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MINERALOGY 2022 Journal of Mineralogical andPetrological Sciences(2022) 117:004 LETTER Geochemical characteristics of an ophiolitic complex from Mt. Tenzan area, Saga Prefecture, northern Kyushu Yusaku TANAKA, Keisuke EsHIMA and Masaaki OwADA Graduate school of Sciences and Technology for Innovation, Yamaguchi University, Yamaguchi 753-8512, Japan The metamorphic complex from the Mt. Tenzan area in northern Kyushu consists mainly of mafic rocks with small amounts of siliceous, calc-silicate, and ultramafic rocks. These lithofacies can be recognized as an ophio- litic complex. Metamorphosed mafic rocks are divided into two types, amphibolites I and II, which are probably derived from supracrustal and intrusive rocks, respectively. The geochemical data of both amphibolites plotted within the field between mid-ocean ridge and island arc basalts; such geochemical features resemble those of back-arc basin basalts. As the metamorphic complex was intruded by Cretaceous granitoids, protoliths of the complex could have been formed prior to the Cretaceous. The protolith lithofacies assemblage and geochemical constraints of the Tenzan metamorphic complex indicate the correlation with the Yakuno ophiolite rather than the Oeyama ophiolite. Keywords: Whole-rock geochemistry, Metamorphosed mafic rocks, Ophiolite complex, Tenzan Area, North- ern Kyushu INTRODUCTION The metamorphic rocks from northern Kyushu occur as large blocks in the Cretaceous granitoids. The protoliths of these metamorphosed rocks are composed of pelitic, calc-silicate, and mafic rocks, with small amounts of lime- stone and ultramafic rocks (Karakida et al., 1969). The Figure 1. (a) Location of northern Kyushu, southwest Japan. (b) zircon uranium-lead (U-Pb) dating of the metamorphosed Regional distribution map of metamorphic rocks from northern clastic rocks from northern Kyushu indicates that the ages Kyushu (modified after Kubo et al., 1993; Hoshizumi et al., of the youngest peaks obtained from detrital zircon cores 2004) with the youngest peak ages obtained from the detrial are approximately 400 and 250 Ma, which correspond to zircon cores from clastic rocks. (c) Simplified geological map the Renge and Suo high-pressure metamorphic belts, re- of the Tenzan area. The zircon U-Pb data are from Tsutsumi et al. (2003, 2011), Adachi et al. (2012), Miyazaki et al. (2017), spectively (Fig. 1; Tsutsumi et al., 2003, 2011; Adachi et and Yuhara et al. (2021). al., 2012; Miyazaki et al., 2017; Yuhara et al., 2021). The Tenzan area in Saga Prefecture is underlain by metamorphosed mafic rocks with small amounts of sili- zan area. Based on the metamorphic pressure-temperature ceous, calc-silicate, and ultramafic rocks (Oshima, 1964) conditions and lithological features, Yamada et al. (2008) (Fig. 1). According to the previous study (Nishimura. correlated the metamorphic rocks from the Tenzan area 1998), the Tenzan area has been thought to be a member with those of the Higo metamorphic rocks without any of the Renge belt because the regional foliations are con- whole-rock analyses. However, the lithological assem- tinued to the high-pressure/low-temperature metamorphic blage of the Tenzan area is equivalent to an ophiolite com- rocks in the Sasaguri area, 40 km north-east from the Ten- plex defined by Ishiwatari (2010), which indicates at least two members of the following lithofacies: 1) mantle peri- doi:10.2465/jmps.210831 dotite, 2) mafic and/or ultramafic cumulate, and 3) mafic M. Owada, owada@ yamaguchi-u.ac.jp Corresponding author Y. Takana, alter8069 @ gmail.com volcanic rocks. This study examined the mineralogical @ 2022 Japan Association of Mineralogical Sciences and whole-rock geochemical characteristics of metamor- 2 Y. Tanaka, K. Eshima and M. Owada phosed mafic and ultramafic rocks from the Tenzan area Amphibolite II is coarse-grained and is composed of and compared our geochemical results with those of other hornblende and plagioclase with small amounts of biotite, Paleozoic ophiolite complexes in southwest Japan. This is potassium feldspar, and quartz, containing ilmenite, titan- the first geochemical analysis of metamorphosed mafic ite, and apatite as accessory minerals (Fig. 2b). Horn- and ultramafic rocks in this area. The data presented here blende rarely developed zonal structure and frequently provide useful information for geotectonic comparisons contains poikilitic plagioclase inclusions. Potassium feld- with Paleozoic ophiolite complexes in southwest Japan. spar appears in the plagioclase grains as an antiperthitic phase. FIELD OCCURRENCE AND PETROLOGICAL Serpentinite is composed of mesh-structured serpen- DESCRIPTIONS tine, talc, tremolite, clinopyroxene, and phlogopite (Fig. 2c). We performed a chemical analysis of the olivine and In terms of field occurrence, the metamorphosed mafic chromian spinel in serpentinite using a JEOL JXA-8230 'I pe I sooque sd on ou ppp are so electron probe micro-analyzer (EPMA) at the Center for Amphibolite I is fine-grained and well-foliated, with thin Instrumental Analyses, Yamaguchi University. Analytical layers and lenses of siliceous and calc-silicate gneisses. procedures of mineral chemistry were performed as de- In contrast, amphibolite II is coarse-grained with weak scribed by Eshima (2021). Olivine (F092-90) appears lo- foliation locally accompanied by serpentinite and tremo- cally as a relict mineral and includes chromian spinel lite rocks as lenticular bodies or lenses (Fig 1c). Based on with a dark brownish color. Its chrome number [Cr# = textural evidence, the mafic and ultramafic rocks in this Cr/(Cr + Al)] ranges from 0.85 to 0.89. The tremolite area underwent a thermal effect at the boundary with the displays columnar shapes with a maximum length of Cretaceous granitoids, but outside of the boundary, the 2.0 cm. The phlogopite exhibits fake crystals surround- rocks were barely affected by contact metamorphism ing the olivine crystals. Mineral data for olivine and chro- (Yamada et al., 2008). mian spinel are available upon request from the corre- Amphibolite I has a fine-grained and nematoblastic sponding authors. texture. It consists mainly of hornblende and plagioclase The tremolite rock primarily consists of tremolite with small amounts of quartz and includes ilmenite, titan- and clinopyroxene with small amounts of serpentine, ite, and apatite as accessory minerals (Fig. 2a). The horn- chlorite, and opaque minerals (Fig. 2d). The tremolite blende has a zonal structure with a light-green core, a is weakly foliated. greenish-brown mantle, and a light-green rim. In addi- tion to these minerals, the calc-silicate thin layers inter- WHOLE-ROCK GEOCHEMISTRY calated with amphibolite I include clinopyroxene, epi- dote, and calcite. The whole-rock chemical compositions of amphibolites I and II, serpentinite, and tremolite rocks were determined by X-ray fuorescence (XRF) analysis. The analyzed samples were free from alteration and veins, and the amounts of 300-600 g were crushed by hand with a W-mortar. Crushed samples were powdered using an au- tomatic W-mortar. Detailed analytical procedures are de- scribed in Eshima and Owada (2018). After ignition in a furnace at 950 °C for 2 h, the samples were measured for loss on ignition; the ignited samples (1.0 g) were mixed with five times the amount of Li2B4O7 as flux. The mixed samples were melted to make glass beads using a bead sampler. The glass beads were analyzed using an XRF analyzer (Rigaku ZSX primus-II) installed at the Center of Instrumental Analysis, Yamaguchi University. The an- 1.0mm alytical conditions included an electric voltage of 50 kV and an electric current of 60 mA, using a Rh anode X-ray Figure 2. Thin section photographs. Plane-polarized light image tube. The analyzed elements were SiO2, TiO2, Al2O3, of (a) amphibolite I, (b) amphibolite I, and (c) serpentinite. Fe2O3, MnO, MgO, CaO, Na2O, K2O, and P2O5 as the Cross-polarized light image of (d) tremolite rock. Hbl, horn- blende; Pl, plagioclase; Ol, olivine; Cpx, clinopyroxene; Tlc, major elements and were Ba, Cr, Nb, Ni, Rb, Sr, V, Y, Zn, talc; Srp, serpentine; Tr, tremolite; Spl, spinel. and Zr as trace elements. To validate the quantitative da- Geochemical characteristics of an ophiolitic complex from Tenzan area 3 Table 1. Chemical compositions of GSJ standard samples, JB-2, Table 2. Representative whole-rock chemical composition of met- JB-3, and JGb-2 amorphosed mafic rocks from the Tenzan area Sample JB-2 JB-3 JGb-2 Type Amp-I Amp-I Amp-I Amp-I Amp-II Amp-II Sample 1903 1903 1903 1903 1905 1905 r.v. m.v. r.v. m.v. r.v. m.v. no. were measured using fused glass beads and a 213 nm 1401 2601 2603 crimination diagram using Nb-Zr-Y (Meschede, 1986) 0302B (wt%) SiO2 52.52 50.77 51.27 47.13 46.73 (wt%) 52.96 SiO2 1.43 0.57 0.55 49.14 50.11 48.09 48.83 50.85 49.28 TiO2 1.18 1.18 1.42 Al2O3 14.56 14.45 17.13 17.09 23.81 23.51 TiO2 1.39 1.58 1.40 1.48 shows a LREE-enriched pattern with an overall 10-350 1.06 Fe2O3 14.17 11.99 Al2O3 15.25 15.06 15.29 15.31 17.33 14.46 11.78 6.78 6.83 18.52 MnO 0.22 0.22 0.18 0.18 0.13 0.13 Fe2O3 9.62 10.46 10.57 10.37 9.70 shows that the IWY-Q2 plot between trachyte to phonolite MnO 4.59 4.57 5.17 5.37 6.27 6.18 0.16 0.17 0.18 0.19 0.16 0.15 MgO 9.77 MgO 7.27 7.44 9.86 9.75 9.76 14.33 8.32 7.29 5.69 5.43 CaO 14.30 Na2O 2.03 1.91 2.72 2.68 0.93 0.92 CaO 13.62 11.82 11.75 12.25 9.39 13.27 K20 Na2O 0.42 0.43 0.78 0.78 0.06 0.06 2.74 Zircon grains were separated from rock powder by 206pb age of 1099.0 ± 0.6 Ma; Paces and Miller, 1993) 2.95 2.97 fields, samples from Akiyoshi-dai plot in basalt field, and P2O5 0.10 0.10 0.29 0.31 0.02 0.03 K2O 0.21 0.20 (with intergrown quartz and Fe-oxide) are anomalous in 0.21 0.80 0.63 100.0099.27 P2O5 0.13 0.17 0.12 0.15 0.10 0.12 Total 100.0099.70 100.00100.85 standard of JEOL was used for the analysis of Hf (HfO2 = 0.44 0.34 0.91 0.75 1.55 (ppm) 1.19 Ba 222 213 245 273 37 60 Total 99.95 100.70 99.62 99.77 99.29 100.62 Cr 28 25 58 63 125 115 (ppm) 2 Ba 45 10 15 59 123 Nb 1 3 3 2 1 131 Ni 17 2 36 31 14 10 Cr 350 254 324 268 100 165 Rb 7 5 15 11 3 3 Ni 77 54 75 83 39 47 IS 180 403 407 438 438 Rb 4 5 4 7 27 178 16 V 575 576 382 V 372 174 174 256 295 271 277 272 212 Y 25 26 27 24 5 Zn 7 67 73 76 73 70 56 ICP-MS Zn 108 110 100 103 49 47 Zr 51 51 98 100 12 13 Sr 255.00 218.00 225.00 274.00 251.00 221.00 Y 26.20 27.60 26.70 25.60 15.10 21.20 r.V., recommended value; m.v., measured value; GSJ, Geological Zr 82.00 94.00 78.00 88.00 41.00 67.00 Survey of Japan. Nb 2.00 15 μm of laser ablation beam diameter. 1.50 3.60 0.40 0.80 La 4.39 4.90 3.53 5.59 5.44 7.90 Ce 11.80 13.80 10.40 14.10 12.40 19.30 ta, we measured the standard rock samples, JB-2, JB-3, Pr 2.01 2.13 1.78 2.28 1.90 2.87 and JGb-2, provided by the Geological Survey of Japan. PN 10.80 11.50 9.74 11.40 8.95 14.30 The results are presented in Table 1. The measured values Sm 3.47 3.75 3.29 3.57 2.47 3.71 of the standard samples were identical to the recommend- Eu 1.23 program Pepi-AGE (Dunkl et al., 2008), and final statis- 1.26 1.27 0.86 1.26 ed values. In addition to XRF analyses, trace elements, Gd 3.89 4.21 3.86 3.80 2.59 3.72 including rare earth elements (REEs), for amphibolites I Tb 0.75 0.77 0.72 0.72 0.45 0.66 and II were determined by inductively coupled plasma Dy 4.91 4.91 4.90 4.74 2.81 3.97 mass spectrometry (ICP-MS) at Activation Laboratory Ho 0.99 1.05 1.03 0.98 0.58 0.80 Ltd., Canada. Er 2.85 2.94 2.99 2.78 1.68 2.36 A total of 34 samples (15 of amphibolite I, 13 of Tm 0.41 0.44 0.43 0.40 0.24 0.35 amphibolite Il, 4 of serpentinites, and 2 of tremolite Yb 2.59 2.83 2.68 2.52 1.48 2.22 rocks) were analyzed by XRF, and 9 samples of amphib- Lu 0.39 0.42 Zircon U-Pb isotope analysis was performed using a 0.38 0.22 0.34 olites I and II were analyzed using ICP-MS. The repre- JH 2.20 2.50 2.10 2.20 1.30 2.00 sentative results are presented in Table 2. All of the Amp-I, amphibolite I; Amp-II, amphibolite II; LOI, loss on igni- whole-rock data used in this study can be requested from tion. the corresponding authors. Figure 3 shows the total alkali versus silica (TAS) and FeO*/MgO versus SiO2 wt% di- agrams (FeO* = 0.9 × Fe2Os*). Amphibolites I and II contents show that the geochemical features of amphib- belong to the subalkaline and tholeite series, respectively olite I are more evolved than those of amphibolite I (Fig. (Figs. 3a and 3b). The FeO*/MgO ratios and Cr (ppm) 3b, Table 2). Y. Tanaka, K. Eshima and M. Owada a Basalt Mg=Mg/ Mg+Fe2)0.2 St Si02 1% Figure 4. Chemical composition of chromian spinel and olivine in SiO2 wt% (b) the serpentinite from the Tenzan area. (a) Compositional rela- TH Cr#. Olivine spinel mantle array (OSMA) is after Arai (1994). CA (b) Compositional relationships between Mg# and Cr# of chro- mian spinel. The data of this study are plotted as the average value from one sample. The compositional ranges of the data are shown as bars accompanied with data plots. The mantle perido- tites from the Oeyama (Arai, 1980; Tsujimori, 1998; Machi and Ishiwatari, 2010) and Yakuno ophiolites (Ishiwatari, 1985a, b) are shown for comparison. Figure 3. (a) SiO2 wt% versus Na2O + K2O wt% (TAS) diagram. (b) FeO*/MgO versus SiO2 wt% diagram (Miyashiro, 1974) those of the Yakuno mantle peridotite (Cr# = 0.6-0.8) showing tholeitic (TH) and calc-alkaline (CA) fields. FeO* = 0.9 × Fe2O3*. than to those of the Oeyama peridotite (Cr# = 0.3-0.5) (Fig. 4). Figure 5a shows the compositional range of primi- GEOTECTONIC SETTING OF TENZAN tive basaltic magmas with their differentiation trends and METAMORPHIC COMPLEX the accumulation directions of specific minerals (Kemp- ton et al., 1997). Kempton et al. (1997) stressed that met- Amphibolite I intercalates with thin layers and lenses of amorphosed mafic rocks up to amphibolite facies grade siliceous and calc-silicate gneisses. In contrast, amphib- should be adopted in this diagram to determine the geo- olite II contacts with lenticular bodies or lenses of ser- chemical compositions of their protoliths. Amphibolites I pentinite and tremolite rocks, and they probably occur as and II studied here underwent amphibolite facies meta- xenoblocks (Fig. 1c). Based on the field occurrence, pro- morphism (Yamada et al., 2008). Therefore, the geo- toliths of amphibolites I and II would, therefore, be su- chemical compositions of the studied samples can be pracrustal and intrusive rocks, respectively. The litholog- adopted to the diagram of Kempton et al. (1997) for es- ical assemblage, including amphibolites I and II and timating the characteristics of the protoliths. The ana- serpentinite in the Tenzan area, is equivalent to an ophio- lyzed samples of amphibolites I and II generally show lite complex as defined by Ishiwatari (2010). Because the evolved compositions, with some samples of amphibolite samples from the Tenzan area were intruded by Creta- I plotted in the primitive basaltic field (Fig. 5a); however, ceous granitoids, they can be recognized as a pre-Creta- the samples do not show accumulation trends. Therefore, ceous ophiolitic complex. Therefore, the metamorphic the geochemical data for amphibolite I and II reflect the complex from the Tenzan area can be compare to geo- liquid composition. The mafic rocks from the Yakuno chemical features of the Oeyama or Yakuno ophiolites. ophiolite also possess the similar compositional ranges, The rock types of the Oeyama ophiolite are dominat- but those from the Oeyama ophiolite are plotted outside ed by mantle peridotite and cumulate rocks (Kurokawa, the liquid compositions. Figure 5b depicts the Mg# ver- 1985). In contrast, the Yakuno ophiolite includes supra- sus TiO2 wt% diagram showing the data from the Tenzan crustal rocks in addition to cumulate rocks and mantle area, the Yakuno and the Oeyama ophiolites (Tsujimori peridotites (Ishiwatari, 1985a). Figure 4 shows the chemi- and Ishiwatari, 2002; Ichiyama and Ishiwatari, 2004; Su- cal composition of chromian spinels and olivines in ser- da and Hayasaka, 2009; Kimura and Hayasaka, 2019). pentinite from the Tenzan area and Oeyama and Yakuno All data were plotted within the field, from mid-ocean ophiolites. Based on the Cr# of chromian spinel and the ridge basalt (MORB) to island arc basalt (IAB). A dia- Fo value of coexisting olivine, the serpentinite in the Ten- gram of the N-MORB-normalized La/Y and Nb/La ratios zan area is a highly depleted mantle peridotite, probably are plotted for amphibolites I and II and the mafic rocks harzburgite or dunite (Fig. 4a). These values are closer to from the Yakuno and Oeyama ophiolites, as well as the Geochemical characteristics of an ophiolitic complex from Tenzan area (b) that the crustal evolution of the Yakuno ophiolite was characterized by intra-oceanic island arc and back-arc basin settings (Figs. 5c and 5d). Amphibolites I and II from the Tenzan area could have experienced the crustal AL evolution in the island arc and back-arc basin settings similar to the Yakuno ophiolite. SiO2 / A1Os (c) ACKNOWLEDGMENTS This work was supported by JSPS KAKENHI Grant Numbers JP15H03748 to M. Owada and JP21J13600 to 0.4 人。 K. Eshima. We wish to thank Dr. K. Itano, two anonymous (La/Y) reviewers, and Dr. Y. Ichiyama, the associate editor. Their comments were useful for improving the manuscript. Rock/C1chondrit (e) REFERENCES Adachi, T., Osanai, Y., Nakano, N. and Owada, M. (2012) LA- ICP-MS U-Pb zircon and FE-EPMA U-Th-Pb monazite dat- <LLD Figure 5. Whole-rock major and trace element compositions of the Mountains, northern Kyushu. Journal of Geological Society <LLD of Japan, 118, 39-52. <LLD Arai, S. (1980) Dunite-Harzburgite-Chromitite complexes as re- Mg# versus TiO2 wt% diagram. (c) N-MORB-normalized La/ fractory residue in the Sangun-Yamaguchi zone, western Ja- Nb versus La/Y diagram (Suda et al., 2014). (d) N-MORB nor- pan. Journal of Petrology, 21, 141-165. malized spider diagram. (e) C1 chondrite-normalized REE pat- Arai, S. 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(2009) Genesis and evolutional process- Released online publication June 4, 2022 es of the Paleozoic oceanic island arc crust, Asago body of the Manuscript handled by Yuji Ichiyama
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Tanaka (2022) - geochem characteristics of an ophiolite Complex Tenzan area, Kyushu.txt
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ELSEVIER Tectonophysics 290 (1998) 197±210
Crustal structure and tectonics of the Hidaka CollisionZone, Hokkaido
(Japan), revealed by vibroseis seismicre¯ection and gravity surveys
KazunoriAritaa,,Takashi Ikawab,T anioItob,AkihikoYamamotoc,MatsuhikoSaitoa,
YasunoriNishidaa,HideyukiSatoha,G a k uK imu r ad,1, Teruo Watanabea, TakeshiIkawae,
ToruKurodae
aDepartment of Earth and Planetary Sciences, Hokkaido University, Kita-ku, Sapporo 060, Japan
bDepartment of Earth Sciences, Chiba University, Inage-ku, Chiba 263, Japan
cResearch Center for Earthquake Prediction, Hokkaido University, Kita-ku, Sapporo 060, Japan
dDepartment of Earth Sciences, CIAS, University of Osaka Prefecture, Sakai-shi, Osaka 593, Japan
eJapex Geoscience Institute Inc., Shinagawa-ku, Tokyo 140, Japan
Received 21 November 1996; accepted 16 December 1997
Abstract
This study is the ®rst integrated geological and geophysical investigation of the Hidaka Collision Zone in southern
Central Hokkaido, Japan, which shows complex collision tectonics with a westward vergence. The Hidaka Collision Zone
consists of the Idon'nappu Belt (IB), the Poroshiri Ophiolite Belt (POB) and the Hidaka Metamorphic Belt (HMB) with the
Hidaka Belt from west to east. The POB (metamorphosed ophiolites) is overthrust by the HMB (steeply eastward-dippingpalaeo-arc crust) along the Hidaka Main Thrust (HMT), and in turn, thrusts over the Idon'nappu Belt (melanges) along
the Hidaka Western Thrust (HWT). Seismic re¯ection and gravity surveys along a 20-km-long traverse across the southern
Hidaka Mountains revealed hitherto unknown crustal structures of the collision zone such as listric thrusts, back thrusts,frontal thrust-and-fold structures, and duplex structures. The main ®ndings are as follows. (1) The HMT, which dips steeply
at the surface, is a listric fault dipping gently at a depth of 7 km beneath the eastern end of the HMB, and cutting across
the lithological boundaries and schistosity of the Hidaka metamorphic rocks. (2) A second re¯ector is detected 1 km belowthe HMT re¯ector. The intervening part between these two re¯ectors is inferred to be the POB, which is only little exposed
at the surface. This inference is supported by the high positive Bouguer anomalies along the Hidaka Mountains. (3) The
shallow portion of the IB at the front of the collision zone has a number of NNE-dipping re¯ectors, indicative of imbricatedfold-and-thrust structures. (4) Subhorizontal re¯ectors at a depth of 14 km are recognized intermittently at both sides of the
seismic pro®le. These re¯ectors may correspond to the velocity boundary (5.9±6.6 km =s) previously obtained from seismic
refraction pro®ling in the northern Hidaka Mountains. (5) These crustal structures as well as the back thrust found in theeastern end of the traverse represent characteristics of collisional tectonics resulting from the two collisional events since
the Early Tertiary. 1998 Elsevier Science B.V. All rights reserved.
Keywords: crustal structure; Hidaka Collision Zone; seismic re¯ection; gravity survey; collision tectonics; Hokkaido
Corresponding author. Tel.: C81 11 706-5305; Fax: C81 11 706-5305; E-mail: arita@cosmos.sci.hokudai.ac.jp
1Present address: Geological Institute, University of Tokyo, Bunkyo-ku, Tokyo 113, Japan.
0040-1951/98/$19.00 1998 Elsevier Science B.V. All rights reserved.
PIIS0040-1951(98)00018-3
198 K. Arita et al./Tectonophysics 290 (1998) 197±210
1. Introduction
The northern island of Japan, Hokkaido, is sit-
uated at the conjunction of two active island
arc±trench systems; the Northeast Honshu Arc±
Japan Trench and the Kuril Arc±Trench (Fig. 1a).Hokkaido is divided into three major provinces with
regardtoitspre-Neogenegeology,thetrendofwhich
is oblique to these active arc±trench systems: West-ern,CentralandEasternHokkaido(Fig.1b).Western
Hokkaido, which is a northern extension of North-
east Honshu, consists of Jurassic accretionary com-
plexes intruded by Cretaceous granitic intrusions.
Fig. 1. (a) Sketch map showing plate tectonic setting around Japanese islands. Thick lines and dotted line show present and Tertiary plate
boundaries, respectively. (b) Geologic divisions of Central Hokkaido. Pro®ling lines Fig. 1cFig. 2a,b, and 3 are shown. Modi®ed fromKimura (1994). (c) Schematic geological cross-section of southern Central Hokkaido. POBDPoroshiri Ophiolite Belt, HMBDHidaka
Metamorphic Belt, HWTDHidaka Western Thrust, HMTDHidaka Main Thrust, TFDTokachi Fault System. Cross-section line is
shown in (b). Modi®ed from Kimura (1986).Eastern Hokkaido is composed of Late Cretaceous
to Palaeogene forearc sediments of the palaeo-Kuril
arc±trench system (Kiminami and Kontani, 1983).
Central Hokkaido, which geologically continues to
Sakhalin, has a complicated geological assemblage.
It has been occupied by two subduction±accretionsystems between the palaeo-Eurasian and palaeo-
North American Plates; the one in the west (the
Sorachi-Yezo Belt, the Idon'nappu Belt and the Hi-daka Belt) is a N±S-trending, westward-subduct-
ing system of Late Jurassic to Palaeogene age, and
the other in the east (the Tokoro Belt) is an east-
ward-subducting system of Cretaceous age (Fig. 1b
K. Arita et al./Tectonophysics 290 (1998) 197±210 199
and c; Kiminami and Kontani, 1983; Sakakibara,
1986).
CentralHokkaido,especiallyitssouthernpart,has
undergone two stages of collisions since the Early
Tertiary. The older oblique collision with a right-
lateral sense of motion between the palaeo-Eurasianandpalaeo-NorthAmericanPlatesduringthePalaeo-
gene resulted in the amalgamation of the Hidaka
MetamorphicBeltintheeastandthePoroshiriOphi-olite Belt in the west. The younger collision was
caused by the westward migration of the Kuril fore-
arc plate which started in the Late Miocene due to
the oblique subduction of the Paci®c Plate along the
Kuril Trench (Fig. 1a; Kimura, 1986). The collisionrapidly uplifted the Hidaka Metamorphic Belt as
well as the Poroshiri Ophiolite Belt, producing the
present Hidaka Mountains. The process of collisionbetween the Eurasian Plate and the North American
Plate(Kurilforearcplate)isstillcontinuing.
Although the deformational features due to the
above two collisions are also observed in the So-
rachi-YezoBelt,the HidakaMountains andtheirsur-roundings in southern Central Hokkaido show typi-
callycomplexcollisionalfeatures.Theyaretherefore
collectively termed as the Hidaka Collision Zone,whichprovidesfavourable opportunitiesforstudying
the characteristicsand structure of deep crust similar
to the Ivrea zone of the Alps and the Kohistan±
Ladakh arc of the western Himalaya. It is impor-
tant to elucidate the present deep-crustal structureof the Hidaka Collision Zone in order to understand
the arc±arc collisionalmechanism and the formation
processofcontinentalcrust.Towardthisend,wecar-ried out vibroseis seismic re¯ection pro®ling along
a 20-km-long traverse, and gravity measurements in
the southern Hidaka Mountains.
2. Geophysicalcharacteristicsof theHidaka
CollisionZone
2.1. Seismicrefraction
Seismic refraction experiments and natural earth-
quakedatainandaroundtheHidakaMountainshave
revealed the following complicated tectonic features
(e.g., Den and Hotta, 1973; Okada et al., 1973;Takanami, 1982; Fujii and Moriya, 1983; Moriya,
1983; Miyamachi and Moriya, 1984; Furumura andMoriya, 1990; Miyamachiet al.,1994; Moriya et al.,
1994; Ozelet al.,1996; Iwasaki etal.,1998);
(1) Beneath the Hidaka Collision Zone there are
two seismic zones: a shallower zone (20 to 50 km
deep) likely related to the collision of the Eurasian
and North American Plates, and a deeper zone (over70 km deep) related to the subduction of the Paci®c
Plate.
(2) In the southern Hidaka Mountains a low-
velocity zone (5.5 km =s) dips eastward from the
western coast, and reaches a depth of about 30 km
beneaththe Hidaka Mountains (Fig. 2a).
(3) The seismic velocity structure of both sides
of the Hidaka Mountains is relatively clear, but thatbeneath the mountains is monotonous or not deter-
minable (Fig. 2).
(4) Lateral variation in seismic velocity due to
different geological units is found across the Hidaka
Collision Zone. In particular, seismic waves are at-
tenuated under the Idon'nappu Belt on the western
part of the Hidaka Mountains.
(5) The Moho discontinuity is not clearly visible
as it may be deeper than 50 km (Fig. 2a), although
Miyamachi et al. (1994) reported a crustal thickness
of 32 km beneaththe Hidaka Mountains.
2.2. Magnetotellurics
A recent magnetotelluric survey (Ogawa et al.,
1994) across the Hidaka Collision Zone (about 70km north of the present survey area: Fig. 3) shows
that a 2-km high-resistivity (1000±2000 /DELm) layer
of high-grade Hidakametamorphicrocks is followedby a relatively conductive (500±1000 /DELm) layer of
accretionary prism at depths of 5 to 10 km, which
is underlain again by a high-resistivity (30,000 /DELm)
layer of probably high-grade metamorphic rocks.
This suggests an inter®ngered complex structure orcrustal delamination beneath the Hidaka Collision
Zone probably due to the above-mentioned two col-
lisionevents.
3. Geologicaloutlineof theHidaka collisionzone
with specialreferenceto the surveyedarea
The Hidaka Collision Zone is occupied by the
N±S-trending accretionary and melange complexes
consisting of the Idon'nappu and Hidaka Belts (eg.,
200 K. Arita et al./Tectonophysics 290 (1998) 197±210
Fig. 2. (a) Seismic velocity structures (km =s) across the southern Hidaka Mountains (after Moriya et al., 1994). (b) Seismic velocity
structures (km =s) across Central Hokkaido revealed by seismic refraction pro®ling (after Ozel et al., 1996). HMTDHidaka Main Thrust.
Both pro®ling lines are shown in Fig. 1b.
Kiminami and Kontani, 1983; Kiyokawa, 1992;
Kimura, 1994). The Hidaka Mountains (the Hi-daka Metamorphic Belt and the Poroshiri Ophio-
lite Belt) are situated between the Hidaka Belt and
the Idon'nappu Belt (Fig. 1b). The Hidaka Colli-sion Zone presently shows westward vergence and a
curvature with a westward convex shape due to the
above-mentionedcollisionof the westward-plunging
Kurilforearcplate.
The Hidaka Metamorphic Belt presently displays
steeply eastward-tilted metamorphic and magmatic
sequences with a general NNW trend, but originallyithadawest-side-downgeometrybefore theamalga-
mationof theHidakamagmaticarcandthePoroshiriophiolites. It is considered to have formed in the
western part of the Hidaka Belt during the oblique
collision of the palaeo-Eurasian and palaeo-NorthAmerican Plates during the Palaeogene (Komatsu
et al., 1983); this view, however, has been debated
(e.g., Kimura et al., 1983; Komatsu et al., 1989;
Maeda,1990; Toyoshima,1991; MaedaandKagami,
1996). According to radiometric dating (Arita et al.,1993), thepresentsteeplyeastward-dippingstructure
of the Hidaka Metamorphic Belt had formed before
K. Arita et al./Tectonophysics 290 (1998) 197±210 201
Fig. 3. Interpreted resistivity structure across the Hidaka Collision Zone (after Ogawa et al., 1994). Cross-section line is shown in Fig. 1b.
the Middle Miocene. The Hidaka Metamorphic Belt
consists of a lower metamorphic sequence (gran-
ulite-faciesrocks and orthopyroxene tonalites)in thewest and an upper metamorphic sequence (biotite±
muscovite gneisses and schists) in the east (Fig. 4).
It decreases gradually in metamorphic grade from
the granulite facies in the west to the greenschist fa-
cies eastward, and in general grades into the weaklymetamorphosedturbiditesoftheNakanogawaGroup
(the Hidaka Belt) of Palaeogene age (Fig. 4: Os-
anai et al., 1986; Komatsu et al., 1994). Thesemetamorphic rocks have been intruded by a large
amount of various intrusive rocks, e.g., gabbroic
anddioriticrocksandS-typeorthopyroxene tonalites
in the lower metamorphic sequence and cordierite
tonalite and granite in the upper metamorphic se-quence (Komatsu et al., 1986; Shimura et al., 1992).
The granulite-facies rocks in the western part are
highly mylonitized near the Hidaka Main Thrust(HMT), along which the Hidaka Metamorphic Belt
overthrusts the Poroshiri Ophiolite Belt to the west.
The mylonites suffered dextral ductile shear defor-
mation under the conditions of greenschist facies
(Arita etal.,1986; Toyoshima, 1991).
In the surveyed area the mylonitized granulites
andtonaliteshavestrongmetamorphicandmyloniticfoliations striking N30ë±50ëW and steeply dipping
to the east. The cordierite tonalite, which occupies
the crestline of the Hidaka Mountains, is massive,but has a weak foliation striking N20ë±40ëW and
dipping steeply eastward on the margin. The upper
metamorphic sequence has the same strike and dip
as those of the tonalite. The Nakanogawa Group
shows a monotonous lithology consisting of slateand shale withsome intercalationsof sandstone. The
turbidites dip steeply east, being repeated by folding
and verticalfaulting.
The Poroshiri Ophiolite Belt is composed of
faultedandtightlyfoldedmetamorphosedophiolites,
the original succession of which has been recon-
structed from a basalt to harzburgite tectonite with
a total thickness of at least 5 km (Miyashita, 1983).TherocksofthePoroshiriOphioliteBeltdisplayalot
ofintenseductiledeformationalfeaturesrepresenting
dextral transpression caused by the oblique collision(Jolivet and Miyashita, 1985; Arita et al., 1986; Arai
and Miyashita, 1994). Although the Poroshiri Ophi-
olite Belt is widely distributed in the northern half
of the Hidaka Mountains, it occurs sporadically as a
narrow zone along the HMT (serpentinite of only 80m wide in the surveyed area), and often is missing
inthe southern half(Fig. 4). The Poroshiri ophiolites
202 K. Arita et al./Tectonophysics 290 (1998) 197±210
Fig. 4. Generalized geological map of the central and southern
Hidaka Mountains. HWTDHidaka Western Thrust, HMTD
Hidaka Main Thrust, ROTDRedatoi±Okada Thrust, NOTD
Nitarachi±Oshorobetsu Thrust, HSZDHoroizumi Shear Zone,
TFDTokachi Fault System, PDMt. Poroshiri-dake. Thick line
is a seismic re¯ection line (Fig. 5).
overthrust the Idon'nappu Belt on the west along the
HidakaWesternThrust(HWT)(Fig.1candFig.4).
The Idon'nappu Belt is divided lithotectonically
into two units by the east-dipping Redatoi-Okada
Thrustassociatedwiththinserpentinitebodies(Uedaet al., 1995), namely the Naizawa Complex in the
west and the Horobetsu-gawa Complex in the east
(Fig. 4). Both complexes are composed of melangeand accretionary sediments. Slaty cleavages in these
complexes generally strike NW±SE, and steeply dip
to the east, although their bedding planes are mostly
west-facing (Ueda et al., 1995). The double colli-
sion has made the Idon'nappu Belt a frontal zoneof the Hidaka Collision Zone showing a dextral
strike-slip duplex structure with apparent westwardvergence especially in the Horobetsu-gawa Com-
plex (Kiyokawa, 1992; Ueda et al., 1995). The
Idon'nappu Belt is in tectonic contact with the Cre-
taceous Yezo Supergroup (the Sorachi±Yezo Belt)
and the Miocene formations on the west along the
east-dipping Nitarachi±Oshorobetsu Thrust (Fig. 4).The Miocene sandstones occur along these faults as
well as withinthe Idon'nappu Belt.
4. Vibroseisseismicre¯ectionpro®ling
Vibroseis seismic re¯ection pro®ling was per-
formed along a 20-km-long traverse on Route 236
across the southern Hidaka Mountains. A standardseismic data processing sequence was used, includ-
ing post-stack coherency ®ltering and ®nite-differ-
ence migrations and depth corrections using a 1-Dvelocitymodel (Table1).
Fig. 5 shows an unmigrated depth section in
which the supposed re¯ection phases are indicated
by arrows with numbers. Each phase is interpreted
as follows (numbers correspond to those inFig. 5).
(1)Theintermittentre¯ectorsareclearlytraceable
fromtheHMTonthesurface(aroundRP150)north-
eastward with an angle of 45ë to a depth of 7 kmbelowtheeasternmarginoftheHidakaMetamorphic
Table 1
Field parameters used in the vibroseis seismic re¯ectionexperiment
Source informationSource type 4 vibrators (Y-2400, MK-IV)Interval 50 mSweep frequency 8±45 HzSweep length 16 sNumber of sweeps 10 sweeps =VP
Sweep mode linear up sweepPhase control ground force locking
Receiver information
Natural frequency 8 HzInterval 25 mNumber of geophones 18 geophones =RP (3 series6 parallel)
Layout 1.4 m interval, linear array
Recording information
Number of channels 240Sample interval 4 msRecord length 24 s (after cross-correlation)Low cut frequency 4 Hz, 18 dB =oct
High cut frequency 90 Hz, 72 dB =oct
K. Arita et al./Tectonophysics 290 (1998) 197±210 203
Fig. 5. Unmigrated depth section across the southern Hidaka Mountains. The interpreted re¯ection phases are shown by arrows numbered
1through8which correspond to numbers in the text. Note a sharp eastward bend of the stacking line at RP 660.
204 K. Arita et al./Tectonophysics 290 (1998) 197±210
Fig. 6. An enlarged migrated depth section showing two listric-shaped re¯ectors corresponding to the Hidaka Main Thrust (top) and
Hidaka Western Thrust (bottom) and a duplex structure between them.
Belt.Thesere¯ectionphases areboundaries between
the complex area in the west and the rather trans-
parent area in the east. After migration, the HMT
re¯ectorsshow a listricgeometry (Fig. 6). The HMTis likely to cut across the foliation of metamorphic
rocks and boundaries of lithofaciesat depth.
(2) The HWT re¯ector is not observable at shal-
low levels. At deeper levels, however, a rather clear
re¯ection phase is recognized about 1 km belowthat of the HMT beneath the eastern ¯ank of the
Hidaka Mountains (between RP 550 and 800), and
interpreted to be the HWT. The layer between thesetwo strong events is considered to be the Poroshiri
Ophiolite Belt, and appears to have a duplex struc-
ture in an enlarged migrated depth section (Fig. 6).
In the unmigrated sectionthe west-dipping re¯ectors
(2
0) look to be traced intermittently from a depth of
6.5 km below RP 550 southwestward, although the
tracesbecomevagueaftermigration.
(3) A back thrust dipping west is found around
thenortheasternmostpartof thetraverse.
(4) There is no clear re¯ectivity in the Hidaka
metamorphic rocks. P-wave velocities in the sub-
surface estimated by processing 240-channel data
obtainedfrom the refractionmethodarealmostiden-tical between the different rock units of the Hidaka
Metamorphic Belt. This suggests poor contrast ofimpedance among the rocks of the Hidaka Metamor-
phic Belt.
(5) Beneath the Idon'nappu Belt some complex
re¯ection phases are recognized at depths of severalkm. These re¯ectors may be suggestive of a frontal
thrust-and-fold structure or tectonic stacking of the
Idon'nappu Belt in front of and below the Hidaka
MetamorphicBelt.
(6) Some subhorizontal re¯ection phases can be
observed at a depth of 14 km beneath the eastern
part of the Hidaka Metamorphic Belt and the Hidaka
Belt. A few similar sub-horizontal re¯ection phasesbecome visible beneath the Idon'nappu Belt after
migration.
(7) Possible short re¯ectors at a depth of 20 km
arelikelytobe lowercrustallamination.
(8) Steep re¯ection planes are observed at depths
of over 11 km beneath the western margin of the
seismic line. After migration, these steep planes
are moved outside the traverse. These planes arepresumed tobe of the Nitarachi±Oshorobetsu Thrust
andfaultsintheSorachi±YezoBelt(Fig.4).
5. Gravity survey
Gravity surveys in and around the Hidaka Moun-
tains (Geographical Survey Institute, 1955; Hagi-
K. Arita et al./Tectonophysics 290 (1998) 197±210 205
Fig. 7. Distribution map of the Bouguer anomaly in the southern part of the Hidaka Mountains with a contour interval of 5 mGal after
terrain correction. A±Bis a gravity and seismic traverse. Thick lines are faults. The broken line is the crestline of the Hidaka Mountains.
HWTDHidaka Western Thrust, HMTDHidaka Main Thrust, NOTDNitarachi±Oshorobetsu Thrust, HSZDHoroizumi Shear Zone, TF
DTokachi Fault System, RDMt. Rakko-dake (1472 m).
rawa, 1967; Miyamachietal., 1987) indicatethat the
Bouguer anomalies along the mountains are highly
positive, reaching up to C140 mGal in the north-
ern part (Maruyama et al., 1991). The high positive
anomaly is considered to be due to ma®c and ultra-ma®c rocks in the Poroshiri Ophiolite Belt and the
Hidaka Metamorphic Belt. This is also supported by
ah i g hV
p=Vsratio of over 1.8 in the western ¯ank
of the Hidaka Mountains (Moriya, 1983). The ®rst
precise gravity measurement was performed along a
seismic line (A±B in Fig. 7, more than 200 stations)
and in the surroundings (more than 300 stations) in
ordertoevaluatecrustalstructuremodelsbeneaththeHidaka Mountains.5.1. Bouguer anomaly
The optimum density for gravity reductions in
the study area was estimated to be 2.6615 g =cm
3
using the method proposed by Murata (1993). The
obtained Bouguer anomaly distribution map of the
study area is shown in Fig. 7. A remarkably high
gravity anomaly belt (about 30 km wide) is lo-cated along the crestline of the Hidaka Mountains.
This positiveanomaly reachesits maximum near the
westernperiphery of thecrestline.It isnoted thatthe
correlationof the Bouguer anomalywith topography
is signi®cantly positive in the Hidaka Mountains,which suggests that little crustal `root' exists be-
neath the mountains. In both the easternand western
206 K. Arita et al./Tectonophysics 290 (1998) 197±210
foothills of the mountains, an abrupt decrease in the
Bouguer anomaly is observed, but the patterns of
anomaliesareasymmetric,andtheBougueranomaly
gradients of both sides are different (Fig. 7). The
eastern abrupt gravity decrease corresponds to that
of the Tokachi Fault System which is an active fault(Research Group for Active Faults of Japan, 1991).
Another abrupt decrease is observable on the west-
ern ¯ank of the mountains including the area alongthe seismic line, although no large fault system is
situated there. Along the seismic line, the positive
gravity anomaly increases abruptly around the HMT
on the surface, and reaches its maximum in the
norther half of the seismic line, and then graduallydecreases toward the eastern margin of the Hidaka
Mountains. At the eastern foot of the mountains,
there is a gravity anomaly trough characterized by astrong negative Bouguer anomaly. It corresponds to
the boundary zone between the Hidaka and Tokoro
Fig. 8. A density structural model along the seismic line A±B in Fig. 7 used for the computation (top) and a comparison of observed and
calculated Bouguer anomalies (bottom) computed after Talwani et al. (1959).Belts,whichisburied by athickpileof LateTertiary
to Quaternarysediments.
5.2. Crustal structure
Basedonthepresentseismicre¯ectionresultsand
thegravityandgeologicalconstraints,asimpleblock
model is constructed for crustal structure along the
seismic line (Fig. 8, upper). The eastward-dippingre¯ector, which is interpreted as the HMT, continues
toadepthof 7km.Beneaththeplane,anarrowlayer
of the Poroshiri Ophiolite Belt is probably situated.
A nearly horizontal re¯ection plane is recognized at
a depth of 14 km. In the model construction, thegravitationaleffectsof thedowngoing slabofthe Pa-
ci®c Plate were ignored because they are considered
to be nearlyuniform on the traversewhich is parallelto the Kuril Trench (Fig. 1a). Fig. 8 (lower) shows
a preliminary crustal model along the seismic line
K. Arita et al./Tectonophysics 290 (1998) 197±210 207
based on Talwani's method (Talwani et al., 1959).
The model Bouguer anomaly ®ts well with the ob-
served gravity especially for the western and eastern
parts of the pro®le, whereas in the central part the
computedgravityshowsasystematicincreasewithin
severalmGal.
6. Discussion:crustalstructure of the Hidaka
CollisionZone
On the basis of seismic refraction data, Den and
Hotta (1973) suggested large-scale thrusting of the
crust of Eastern Hokkaido over Western Hokkaido
and vast sedimentation in the foredeep west of thethrust. They attributed the thrust boundary to a plate
boundary between the Okhotsk (the North Ameri-
can) and Eurasian Plates during the Mesozoic. Sucha tectonic scheme around the Hidaka Mountains
(Fig. 2a) has been supported also by recent geophys-
ical studies (e.g., Takanami, 1982; Miyamachi and
Moriya, 1984; Miyamachi etal.,1994; Moriya et al.,
1994).
Fig. 9. An interpreted crustal model of the southern part of the Hidaka Collision Zone. HWTDHidaka Western Thrust, HMTDHidaka
Main Thrust.Although the present integrated study of geologi-
caland geophysicalwork could not detectthe Moho,
it could image the collision tectonics beneath the
Hidaka Mountains such as the listric-shapedHMT, a
back thrust, a fold-and-thrust structure and a subhor-
izontalre¯ectorat a depth of 14 km (Fig. 9).
Asalreadystated,theHidakaMetamorphicBeltis
considered to be an upthrust magmatic arc (palaeo-
Hidaka arc) tilting steeply eastward similar to theHMT on the surface, and therefore the deep rocks of
the palaeo-Hidaka arc occur in its western part. The
granulite-faciesrocksoutcroppinginthewesternpart
of the study area are intensely mylonitized, but, in
general, the thermobarometric analyses of the gran-ulite-facies rocks indicate pressure and temperature
conditions corresponding to those at a depth of 23
km (Osanai et al., 1986). This thickness is almostthe same as the total thickness of the reconstructed
crustal sequences of the Hidaka arc including the
Nakanogawa Group (Komatsu et al., 1983). There-
fore, the depth of the HMT beneath the Hidaka Belt
should be expectedtobe more than23km, assuming
208 K. Arita et al./Tectonophysics 290 (1998) 197±210
that the observed parallel relationship on the surface
between the HMT plane and the lithologic bound-
ary and foliation planes of the Hidaka metamorphic
rocks is maintained at the deeper levels.The seismic
re¯ection pro®ling, however, proves the HMT to be
a listric fault, and by far shallower than that esti-mated before. Hence the HMT most probably cuts
across the lithological boundaries and foliations at
depth (Fig. 9). The Hidaka Main Thrust sheet (Hi-daka Metamorphic Belt) is found to be much thinner
than it has been generally expected. It is noted that
the HMT seems to continue eastward to the velocity
boundary between the 5.9±6.0 km =s layer and the
6.2±6.3 km=s layerdeduced from theseismicrefrac-
tion pro®ling beneath the Tokachi Plain in the north
(Fig. 2b: Ozel et al., 1996; Iwasaki et al., 1998), al-
though both areas are about 60 km from each other.The HMT is the most signi®cant tectonic feature
traceable along the whole Hidaka Mountains, and is
thought to have been a plate boundary until the Ter-
tiary.However,theHMTisjustalistricfaultsituated
at the middle of the upper crust, and consequentlythe true palaeo-plate boundary is expected to exist
beneaththe HMT atdepth.
In the northern half of the Hidaka Mountains, the
Poroshiri Ophiolite Belt reaches up to 5 km in width
(Miyashita, 1983), and large gabbroic bodies occur
in the western part of the Hidaka Metamorphic Belt
(Fig. 4). These ma®c rocks are attributed to the high
positive Bouguer anomaly of up to 140 mGal. Onthe other hand, in the southern half only small bod-
ies of the ophiolitic rocks occur intermittently along
the HMT like in the study area. Nevertheless, thepositive Bouguer anomaly is still high in the present
area (Fig. 7). A 1-km-thick Poroshiri Ophiolite Belt
detected at depth by seismic re¯ection may be re-
sponsible for the positive Bouguer anomaly. This
may be supported by the existence of a conductivelayer between resistive layers beneath the northern
Hidaka Mountains (Fig. 3), although their depth and
thickness are different from each other. The differ-ence in the Bouguer anomaly gradients between the
eastern and western sides of the Hidaka Mountains
and the rapidincreaseof the anomaly from the HMT
eastward (Fig. 7) also suggest that the Poroshiri
ophiolitesgentlydip eastwardand become thickertotheeast.Itisinterestingtonotethatseismicpro®ling
is suggestive of a duplex structure betweenthe HMTand HWTre¯ectors (Fig.6). Theserpentinizedophi-
oliticlayer is considered to play an important role as
a mechanical ¯ow plane for the westward thrusting
of the Hidaka metamorphic rocks along the HMT.
In the unmigrated section the HWT re¯ector is dis-
tinct beneath the northeastern part of the pro®le, andlooks to branch off downward around RP 550, being
traceableintermittentlysouthwestward (2
0in Fig. 5).
If the re¯ection phase is true, this may imply an ex-istence of a tectonic wedge of Indon'nappu melange
which splits the upper crust of the Hidaka arc into
two parts. Further detailed analyses are required to
ascertainthe southwest-dipping re¯ector.
In the shallow part of the Indon'nappu Belt, a
number of NNE-dipping re¯ectors appears, indica-
tive of a fold-and-thrust structure. Ueda et al. (1995)
reportedaduplexstructurewitha right-lateralstrike-slip sense of motion having resulted from dextral
transpression between the Hidaka arc crust and
an oceanic crust (Poroshiri ophiolites) in the Late
Oligocene to Early Miocene (Arita et al., 1986). The
duplex structure probably evolves into the fold-and-thrust structure atdepth.
A west-dipping re¯ector is seen clearly at the
northeastern end of the seismic traverse. On the sur-face some faults are observed, but the sense of their
movement is not clear because of the monotonous
lithology of the Nakanogawa Group. The re¯ector,
however,is presumedtobe abackthrust on thebasis
of a generaltectonic¯ame.
The very weak subhorizontal re¯ectors at 14
km depth at both sides of the seismic pro®le are
signi®cant (Fig. 9). Such a 14-km-deep subhorizon-tal boundary was also detected as a clear velocity
boundary between the 5.9 km =s layer and the 6.6±
6.7 km=s layer by seismic refraction pro®ling in
the northern Hidaka Mountains (Fig. 2b: Ozel et
al., 1996; Iwasaki et al., 1998). These layers aresupposed to be the upper and lower crusts, respec-
tively.In thenorthern HidakaMountains, thedistinct
boundary is traceablewidely from the Sorachi±YezoBelt through the Hidaka Mountains, and dips gen-
tly eastward to the Tokachi Plain (Fig. 2b). Beneath
the Tokachi Plain a middle velocity layer (6.2±6.3
km=s) exists between the layers of 5.9 km and 6.6
km=s in the northern Hidaka Mountains (Fig. 2b).
The boundary between the layers of 5.9 km =sa n d
6.2±6.3 km=s is located at about 8 km depth. It is
K. Arita et al./Tectonophysics 290 (1998) 197±210 209
worth to note that the boundary seems to be the
easterncontinuation of the HMT, although these two
arefarfrom eachother.
Further seismic pro®ling is required eastward for
a detailed imaging of the Hidaka crustal structure
and westward for understanding the tectonics in thecollisionalfront.
Acknowledgements
We would like to thank R. Sorkhabi for criti-
cal reading and improvements of an earlier draft
of the manuscript. This paper has been greatly im-
proved through the efforts of S. Klemperer and twoanomymous reviewers. N. Oshima, H. Yokota and
K. Kameda helped us with the seismicline measure-
ment. Thanks to Y. Murata for providing a computerprogram for calculatingthe Bouguer anomaly by the
ABI method. Financial support for this work was
provided by a grant from the Grant-in-Aidfor Scien-
ti®c Research of the Ministry of Education, Science,
Sports and Culture,Japan (06402018) toK.A.
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Arita (1998) - Crustal structure and tectonics of the Hidaka Collision Zone.txt
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Geochemical Journal. Vol. 18, pp. 195 to 202,1984
Origin
steamsof gases and chemical equilibrium
from Matsukawa geothermal area,among them in
Northeast Japan
YUTAKA YOSHIDA
Geothermal Development Division, Japan Metals and Chemical Co., Ltd.
24 Ukai, Takizawa-mura, Iwate-gun, Iwate 020-01, Japan
(Received July 19, 1983: Accepted June 5, 1984)
Gas components contained in geothermal steams discharged from wells at the Matsukawa geothermal
areas were examined geochemically. The original deep seated gases of Northeast Japan are suggested to be
uniform with respect to He, Ar and N2 and are emitted through geothermal wells and/or fumaroles after
mixing in various proportions with atmospheric air dissolved in ground water.
Geothermal wells of the Matsukawa area are divided into two groups by the geological structure, of
the area which controls the variation in concentrations of tritium and major gas components occurs. The
influence of the geological barrier can be considered to be limited in a shallow horizon.
The correlation between gas components indicates that the reaction, 2NH3 = N2 + 3H2, is in equi
librium, but the reaction, CH4 + 2H20 = C02 + 4H2, is not in equilibrium under the condition of the
Matsukawa geothermal reservoir.
INTRODUCTION
The Matsukawa geothermal area is a vapor
dominated type geothermal system which is
unique among geothermal systems so far ex
plored in Japan. The first geothermal power
station in Japan was completed in this area by
Japan Metals and Chemicals Co., Ltd. in 1966.
At present, 22MW of electricity is generated
by geothermal steam discharged from eight
production wells. As the geothermal steam is
produced directly from zones deeper than
1,000m through a casing pipe, the contamina
tion of organic material, shallow meteoric water
and atmospheric air does not occur significantly
during the passage of steam through the well.
The components of high-temperature vol
canic gases have been studied from the point of
view of chemical equilibrium (ELLIS, 1957;
MATSUO, 1960; HEALD et al., 1963; STOIBER et
al., 1974; GERLACH, 1979). In recent years,
geothermal development has become active all
over the world, and gochemical studies of
geothermal systems have made a great progress
(D'AMORE et al., 1980; GIGGENBACH, 1980).Moreover, rare gases in volcanic gases are studied
recently in relation to the origin of gases
(MATSUO et al., 1978; NAGAO et al., 1980;
TORGERSEN et al., 1982; Kn'osu, 1983a).
In the present study, an attempt is made to
examine the origin of gases on the basis of the
analytical results of steam discharged from
geothermal wells in the Matsukawa area. In this
paper, steam means the mixture of water vapor
and other gas components such as C02, H2S,
H2, He and so forth.
GEOLOGICAL SETTING
The Matsukawa geothermal power station is
located in the Hachimantai volcanic region
which is one of the most active geothermal areas
in Northeast Japan (Fig. 1). Geological investi
gations in this area have been carried out by
NAKAMURA et al. (1961) and SuMI (1966, 1968),
and abundant geological data have been accumu
lated by Japan Metals and Chemicals Co., Ltd.
The basement of this region consists of the
Miocene Kunimitoge, Takinoue-onsen and Ya
matsuda formations, which are composed of
195
196 Y. YOSHIDA
shale, sandstone, tuff and conglomerate, and is
overlain by the Pliocene andesitic-dacitic Tama
gawa welded tuff. The Tamagawa welded tuff is
covered by the Pleistocene Matsukawa andesite
which is regarded as the' cap rock of the geother
mal system in this area. Several Pleistocene
formations composed of andesitic volcanic rocks
overlie the Matsukawa andesite. Most of the
geothermal steams are derived from the lower
part of the Tamagawa welded tuff formation
and the Yamatsuda formation.
The depth of eight production wells and one
exploration well ranges from 1,000 to 1,600m,
and at present, the steam is perfectly dry and
superheated by 20 to 70°C as compared with
the liquid vapor equilibrium temperature at the
well-head pressure.
The altered rock zone extends along the
Matsukawa river in the direction from ENE to
WSW, comprising an area 7km long and 0.5
1.0km wide. The altered rock zone is calssified
into some subzones as shown in Fig. 1, on the
basis of the mode of occurrence of mineralassemblage. These zones are not only distribut
ed horizontally but also vertically as shown in
Fig. 2, and kaolinite, anhydrite, pyrite and
other alteration minerals are found in boring
core and cutting samples.
SAMPLING AND ANALYSES
Sample collection and chemical analyses of
the geothermal steam from wells of the Matsu
kawa geothermal power station were carried
out in September and December 1982, and
steam condensates were analyzed in April 1982.
Samples of steam condensates for the measure
ment of tritium concentration were collected in
1975 and 1980. Localities of wells are shown in
Fig. 1.
The methods of sampling and analyses
were similar to that of OZAWA (1967), but
partly modified for the sake of convenience.
Samples of steam condensates were collected
through a glass coil condenser. Non-absorbable
gases in alkaline solution were analyzed by the
/
///// / 4atsukawa R. 1_4
/ e~@To Q ® y /'8/
// / /
/ / /
// //AA
/M
/y•// /11\/,M1
16 l/
%4kagaa is R.\
(\ /M9
f//
V/
/ /
7 /
/ / Legendloom
sG~
//Zone of
weak alteration
Zone of
montmorillonite
Zone of
kaolinite
Zone of
alunite
E
SEA OF
JAPAN147E
O~INUAA
KAKKONDA1N
PACIFIC
OCEAN
-4ON
39N
Fig. 1.
zones.Map of the Matsukawa geothermal area showing the localities of wells and distribution of altered rock
•: well-head locality, -*: well-bottom locality .
Origin of gases and chemical equilibrium among them 197
gas chromatographic method (SuGisAKI et al.,
1980; KAWABE et al., 1981). The Hitachi model
164 gas chromatograph combined with a pre
amplifier, Ohkura model AM 1001 B micro-volt
meter was used for He, Ar and N2 measure
ments. Tank oxygen was used as carrier gas at
the flow rate of 5 ml/min. The separation
column consisted of teflon tubing (3 mm inner
diameter) 5 m in length packed by 60/80 mesh
Molecular Sieve 5A. The oven temperature was
set at 40'C. Since this gas chromatograph was
modified to remove hydrogen gas by heated
stainless steel column packed with CuO grains,
the Hitachi model 163 gas chromatograph was
also used for H2, N2 and CH4 measurements.
Tank argon was used as carrier gas at the flow
rate of 30ml/min. The separation column
with 3mm (I.D.) stainless steel tubing 2m in
m 800
400
Sea
bevel
-400LN
NQ
HN
m
()It
cc
2length was packed with 60/80 mesh Molecular
Sieve 5A. The oven temperature was set at
40'C. The analytical error for He was about
10% and for the others less than 5%.
The NH3 concentration in steam condensate
was analyzed colorimetrically. The measure
ment of tritium concentrations in steam con
densates was performed at Gakushuin University
by gas counting method described by YONEDA et
al. (1967), after electrolytic enrichment of
tritium.
®I a3 ®5
®2 ®4
Fig. 2. Schematic cross section of alteration zones in
the Matsukawa geothermal field. 1: Zone of montmoril
lonite and iron-rich saponite, 2: Zone of chlorite, 3:
Zone of kaolinite, 4: Zone of alunite, 5: Zone of
pyrophyrite (after KIMBARA, 1983). RESULTS AND DISCUSSION
He, Ar and N2 concentrations in geothermal
steam Results of chemical analyses of
geothermal steam are listed in Table 1. In this
table, all the gas concentrations are expressed
by volume concentration in the steam. The He/
Ar and N2/Ar ratios are further normalized to
the corresponding atmospheric ratios (Fig. 3).
As seen in Fig. 2, sample points are distributed
mostly along the curve which connects point A
with B. Point A shows the dissolved air in water
which is in equilibrium with the atmospheric
air at 10° C and B indicates the gas (M-3, Decem
ber 13, 1982) with the highest ratios of He/Ar
and N2/Ar. Point B can be considered to re
present the gas derived from a deeper horizon of
this area. The curve is a calculated mixing line
of gases with the composition of points A and
B.
In such a small geothermal area as Matsuka
wa, it can be postulated that the deep seated
gas with a homogeneous composition exists
in a deeper horizon of the area. In other words,
variation of He/Ar and H2/Ar ratios of geother
mal steam is not due to the variation in He/Ar
and N2/Ar ratios of the original deep seated
gases but to the change in mixing ratio of deep
seated gases and dissolved air. The mixing may
occur in the reservoir formation processses.
Corresponding ratios of fumarolic gases from
volcanoes of Northeast Japan (Kiyosu, 1983a)
are also distributed along the extended curve
in Fig. 2. The positive correlation between He/
Ar and N2/Ar ratios common for the original
198 Y. YOSHIDA
Table 1. Composition of geothermal steam from geothermal wells at Matsukawa
Gas concentration in steam (by volume)
Well Depth
mDate Total gas H2 S C02 H2 N2 CH4 Ax He NH3
%ppm ppm ppm ppm ppm ppb ppb ppm
M-1
M-2
M-3
M-5
M-6
M-7
M-8
M-9
T-241006
1080
1170
1190
1203
1280
1406
1599
10509/28,1982
12/13,1982
9/28,1982
12/13,1982
9/28,1982
12/13,1982
9/28,1982
12/13,1982
9/28,1982
12/13,1982
9/28,1982
12/13,1982
9/28,1982
12/13,1982
9/28,1982
12/13,1982
12/13,19820.87
0.83
0.30
0.33
0.71
0.74
0.35
0.33
0.32
0.31
0.26
0.24
0.42
0.36
1.14
1.08
0.43487
540
441
492
5 25
555
410
360
368
474
403
384
760
673
616
572
4438010
7550
2480
2730
6350
6660
2960
2810
2630
2510
2110
1910
3360
2830
10600
10100
382022.4
33.0
42.6
44.9
88.0
94.0
32.0
35.0
33.0
55.5
34.6
45.1
30.5
37.1
128
108
12.8115
111
25.9
20.9
88.0
58.1
61.3
65.0
51.2
44.0
36.4
43.2
35.4
38.5
40.2
35.6
19.661.9
64.2
12.2
9.77
50.7
32.5
32.3
34.7
20.7
17.6
14.3
16.7
17.7
17.6
36.7
29.2
6.021220
1120
324
268
930
528
693
726
643
543
523
562
496
479
456
455
28824.0
21.7
3.90
5.02
16.8
14.2
9.21
15.4
8.77
9.05
5.15
7.49
7.48
5.90
6.73
8.42
2.8410.0
15.3
54.4
4.7
5.9
6.2
8.3
51.0
Total gas: Gases other than water vapor.
NH3: Samples were collected on April 20, 1982.
deep seated gas of the Matsukawa area and
fumarolic gases from volcanoes of Northeast
Japan may be related to the fact that the origi
nal deep seated gas of Northeast Japan is uni
form with respect to nitrogen and noble gases
and is emitted through geothermal wells and/
or fumaroles after mixing with dissolved air in
various proportions.
As seen in Fig. 3, He/Ar and N2/Ar ratios of
the original deep seated gas of Northeast Japan
are more than hundred times and three times as
large as those of atmospheric air, respectively.
According to SUGISAKI et al. (1978), four ex
pected sources of nitrogen are as follows: 1.
penetrating atmospheric air, 2. bacterial decom
position of organic matter contained in sedi
ments, 3. pyrolysis of organic matter, 4. release
of inorganic nitrogen from igneous and/or
metamorphic rocks. The most probable origins
of nitrogen which raises the N2/Ar ratio are (3)
and (4) among four possibilities described
above.
On the other hand, MATSUO et al. (1978)
suggested that one of the reasons for high
N2/Ar ratios in volcanic gases from island arcvolcanoes is the contribution of factor (3)
due to sedimentary materials transferred into
the lower crust or upper mantle through subduc
tion.
It can be suggested that N2/Ar ratio of the
original deep seated gas to which there is no
contribution of dissolved air has a fixed value
controlled by sedimentary materials. It can
also be concluded that He, Ar and N2 gases
contained in both of geothermal steams and
fumarolic gases of Northeast Japan are derived
from a common source.
Geothermal reservoir Eight production wells
and one exploration well produce steam at the
Matsukawa geothermal power station. Wells
except M-2 and M-5 were drilled by directional
drilling as indicated in Fig. 1. Table 1 shows
that CO2 concentrations of steam from M-1, 3
and 9 range from 6,350 to 10,600ppm and
those of the others from 1,910 to 3,820ppm.
The H2S concentration of steam has a range
from 360 to 760ppm, and the variation in the
concentration of H2S is smaller than that of
CO2.
Origin of gases and chemical equilibrium among them 199
The relationship among C02, H2S and R
gas (residual gases after the gas is washed with
5 N KOH solution) is shown in Fig. 4. As shown
in Fig. 3, eight production wells are divided
into two groups, i.e., wells M-1, 3 and 9 and
wells M-2, 5, 6, 7, and 8. The exploration well
T-24 does not belong to both of the two groups.
Two production well groups can be distin
guished from each other also by their localities.
Wells M-l, 3 and 9 are located in the zone of
weak alteration, and other wells are located in
the zones with the occurrence of montmoril
lonite, kaolinite and alunite as shown in Fig. 1.
The zonal distribution on the exposed surface
is observed in the vertical section; the alunite,
zone occurs in center, successively surrounded0s
/_Q/_6
00 0°
Ao
00
q)
0°°0
°
>o
1000
100
10
a1.0
0.10
00Q
a
Aeo
0
+B
o ~ 0
00
0 00
0
0.5 1.0 1.5 2.0
( NZ/Ar)sample/(N2/Ar)air
Fig. 3. Relationship between HelAr and N2/Ar rarios of
geothermal gases. •: M-1, 3 and 9, 0: M-2, 5, 6, 7 and
8, o: Fumarolic gas of Northeast Japan (after Kiyosu,
1983a). Point A shows dissolved air in water which is in
equilibrium with atmospheric air at 10°C, and point B
the gas from M-3 with the highest ratios of He/Ar and
N2/Ar. The curve is calculated mixing line of gases of
points A and B. Three fumarolic gases with (N2/
Ar)samplel(N2/Ar)air. > 2 are excluded from this Figure.100 90 80
C02 <
Fig. 4. Gas composition of geothermal steam of Matsu
kawa shown by triangle diagram for C02, H2S and R -gas
(residual gases). •: M-1, 3 and 9, 0: M-2, 5, 6, 7 and 8,
o: T-24 .
by the kaolinite zone and montmorillonite zone
(Fig. 2). Calcite exists in the zones of weak
alteration (SuMI, 1968), and the coincidence of
the localities of wells which discharge CO, rich
steam and the zone of weak alteration suggests
that calcite is one of the sources of CO2 in
steam from wells M-1, 3 and 9.
The tritium concentrations of condensates
of wells M-3 and M-9 are 1.09 1.19 T.U., and
those of wells M-5 and M-8 are 0.28 0.40T.U.
(Table 2).
The variety of the concentrations of major
gas components and tritium as well as He/Ar and
N2/Ar ratios suggests a possibility that the
reservoir or steam channel for wells M-1, 3 and 9
is separated by some barrier (faults, which are
assumed from the upheaval structure of the
Yamatsuda formation) from the other reservoir
Table2. Tritium concentration in condensate of
steam from Matsukawa
Well Date T (T.U.)
M-3
M-5
M-8
M-98/13, 1980
12/ 8.1975
8/13,1980
8/13,19801.09
0.28
0.40
1.19
T. U. = (T/1H) X 1018.
200 Y. YOSHIDA
Table 3. Correlation coefficients(r) between gas components in geothermal steam atMatsukawa
H2S
0.443C02
0.999
0.416 H2
0.682
0.336
0.681N2
0.402
-0.051
0.400
-0.070CH4
0.653
0.107
0.650
0.195
0.944Ar
0.335
-0.083
0.333
-0.147
0.985
0.907He
0.448
-0.009
0.445
-0.005
0.967
0.928
0.935NH3
0.737
0.372
0.735
0.942
0.014
0.275
-0.097
0.094Total gas
H2S
CO2
H2
N2
CH4
Ar
He
or steam channel for wells M-2, 5, 6, 7, 8 and T In this connection, an independent behavior of
24. Since He/Ar and N2/Ar ratios of the original H2S seems to be due to the buffer reaction, e.g.,
deep seated gas of the Matsukawa area are uni sulfide + H20= oxide + H2S . KlYosu (1983b)
form, the influence which causes the variation showed also the possibility that the hydrogen
of concentrations of tritium and chemical com isotopic exchange equilibrium is also established
ponents can be considered to be limited in a between H20 and H2 through the following
shallower horizon than that in which the origi reaction,
nal geothermal gas of the Matsukawa area exists.
2H20= 2H2 +02 (1)
Correlation between individual gas compo
nen is Gas composition of geothermal steam It can be said that weak correlations indicated
gives us two interesting problems. One is the for the pairs of gas components including H2 and
origin of each component and the other is the H2S come from high reactivity of H2 and H2S
chemical reactions taking place in the geother through chemical reactions. Some of possible
mal systems. The origin of He, Ar and N2 has reactions are described above. Beside reaction
been discussed in the previous section. In order (1), following two reactions including H2 can be
to elucidate the origin of gas components such considered among major components to investi
as C02, CH4, H2S, H2 etc., correlation coef gate the reactivity of H2
) ficients(r) between the gas components are cal
culated first, and are given in Table 3. The 2NH3 = N2 + 3H2 (2)
Irl values higher than 0.693 indicate significant
correlation at the confidence level of 99.9%. CH4 + 2H20 = CO2 + 4H2 (3)
Very strong positive correlation (r=0.999)
between C02 and total gas* implies that contri When reaction (2) is in equilibrium, its
bution of CO2 concentration and its variation to equilibrium expression in terms of fugacities fi
those of total gas is great. Strong positive cor becomes
relations are also indicated for the pairs of 3
N2-CH4, N2-Ar, N2-He, CH4-Ar, CH4-He and Ar KN = f (4)
He. N H 3
On the other hand, H2S and H2 have no cor where K
N denotes the equilibrium constant. relation with other gas components, while H2 Since fugacity coefficients can be assumed
shows an appreciabe correlation with NH3. to be close to unity under the condition of KiYosu (1980) showed that the sulfur isotopic common geothermal reservoir, reaction (4) can exchange equilibrium is established at the well be converted using partial pressure, pi, to be bottom temperature between H2S, pyrite and 3
anhydrite in the Matsukawa pN2 PH2 (5) geothermal area . KN = 2
PNH3
* Total gas is defined in this study as all gases except water vapor (refer to Table 1).
Origin of gases and chemical equilibrium among them 201
As the partial pressure can not be measured
directly at present, we have to use the measured
concentrations for discussion. The fact that
neighboring Kakkonda and Ohnuma geother
mal fields (Fig. 1) are water dominated-type
systems suggests that water exists originally in
liquid phase even in the Matsukawa reservoir.
Since the steam is dry, we can assume that a
part of liquid phase in the reservoir evaporates
completely, so that all the chemical components
found at the well-head are originally dissolved in
the liquid. Then, the concentration C; of com
ponent i measured at the well-head can be used
to check the establishment of chemical equi
librium in the liquid reservoir..
Equation (5) can be converted using con
centration of species i in the discharging steam
to equation (6),
CNH3 = A •CN2 • CH2 (6)
where A is a constant at a fixed temperature.
If reaction (2) is in equilibrium, CNH3 should be
proportional to CN2 • CH2. In order to examine
reaction (2), correlation coefficient was cal
culated. A strong positive correlation (r=0.986)
between CNH3 and CN2 • 'CH2 suggests that reaction (2) is in equilibrium.
As the measured well-bottom temperature
at Matsukawa shows a rather narrow range from
223 to 2600C, reservoir temperature is assumed
probably to be constant near 300'C. Rates of
chemical reaction and carbon isotopic exchange
in reaction (3) are considered to be very slow at
temperatures around 300°C (HULSTON, 1977;
SACKETT et al., 1979; NUTI et al., 1980;
GIGGENBACH, 1982). If reaction (3) is in equi
librium at a constant temperature, following
two equations should hold,
2 1 aPCH4 • PH2O = K~ • pco2 ' P H2 (7)
and CCH4 = B • CCO2 • CH2 (8)
where K, and B are an equilibrium constant and
a constant, respectively. In the case of constant
temperature, PH2o (saturated water vapor pressure) can be regarded as constant. Again, the
correlation between CH4 and CC02 • CH2 is examined. The correlation coefficient is ob
tained to be 0.149, which suggests that reaction
(3) is not in equilibrium. It suggests that carbon
isotope exchange reaction between CO2 and
CH4 is also not in equilibrium. This disequi
librium state can be regarded as the result of
contamination of CO2 and/or CH4 from outside
the geothermal system. The contributions of
organic matter to C02 and CH4 production can
not be neglected in some cases (D'AMORE et al.,
1977; WELHAN et al., 1979). It is necessary to
reconsider the applicability of carbon isotopic
geothermometer method.
On the basis of the facts mentioned above,
NH3 is considered to be controlled by chemical
reaction with H2 and N2. It is considered also
that H2 is controlled by chemical reactions
between Fe(II) in minerals, H2O and H2S in the
Matsukawa geothermal reservoir.
Acknowledgement-The author wishes to express his
thanks to Professor N. NAKAI of Nagoya University and
Professor S. MATSUO of Tokyo Institute of Technology
for their critical reading the manuscript. He also ac
nowledges Dr. H. NAKAMURA of Japan Metals and
Chemicals Co., Ltd. for permission and encouragment
which led to the preparation of this paper. Thanks are
also due to the staff member of the geochemical section
of J.M.C. for their help during the sampling and
analyses.
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Journal of Mineralogical and Petrological Sciences, Volume 113, page 219-231, 2018 Compositional variation of olivine related to high-temperature serpentinization of peridotites: Evidence from the Oeyama ophiolite Toshio NoZAKA Department of Earth Sciences, Okayama University, Okayama 700-8530, Japan Compositional variation of olivine in serpentinized peridotites provides a significant constraint on modeling the redox conditions of serpentinization and the tectonothermal history of ophiolites. Here I report the variations of Fe, Mg, Mn, and Ni contents of olivine from the Oeyama ophiolite, SW Japan and show textural and chemical evidence for compositional modification of olivine related to high-temperature (T) serpentinization. The Fe- enrichment of olivine adjacent to antigorite without significant magnetite formation indicates a reducing con- dition for high-T serpentinization. Systematic variations of forsterite (Fo) component with distance from anti- gorite suggest Mg-Fe volume diffusion took place in olivine porphyroclasts under the conditions of high-T diagram, which could be a useful indicator of high-T serpentinization. Retrograde antigorite is different from prograde antigorite in having a shape of elongated blade, lacking a significant amount of magnetite inclusion, and being more ferrous than lizardite. The existence of retrograde antigorite provides another piece of evidence for high-T serpentinization even if olivine has been decomposed by intense low-T serpentinization. Approx- imate estimation of time required for the observed Mg-Fe diffusion profiles of olivine porphyroclasts reveals within the serpentinized mantle wedge following rapid exhumation immediately after the amphibolite-facies metasomatism. Volume diffusion INTRODUCTION Magnetite is commonly associated with lizardite and/or brucite in serpentinites. Recent studies have shown Because of the possible relation to the origin of life and the uneven distribution of magnetite in serpentinites, sug- the development of biosphere, molecular hydrogen pro- gesting a local variation in oxygen fugacity, silica activity duced by serpentinization of peridotites have attracted sig- and/or water-rock ratio (e.g., Bach et al., 2006; Beard nificant interests (e.g., Kelley et al., 2001; Fruh-Green et et al., 2009; Katayama et al., 2010; Frost et al., 2013; al., 2003; Sleep et al., 2004; Takai et al., 2004; Kelley et Miyoshi et al., 2014; Schwarzenbach et al., 2016). In con- al., 2005; Schulte et al., 2006; Russell et al., 2010). The trast, Evans (2010) has proposed that magnetite is absent hydrogen production is mainly caused by the oxidation of at relatively high-temperature (T) antigorite-forming ser- Fe in olivine and associated with the formation of magnet- pentinization, because Fe is distributed in olivine with a ite (e.g., Frost, 1985; McCollom and Bach, 2009; Klein considerable Mg-Fe interdiffusion rate, instead of forming et al., 2013, 2015). In this context, it is important to under- magnetite. Such a behavior of Fe during high-T serpentin- stand the formation conditions of magnetite and the effect ization could be a constraint on hydrogen production at of its presence or absence on serpentinization processes. a mantle wedge of supra-subduction zone, at a deep level of oceanic lithosphere, or at the ancient Earth's surface. doi:10.2465/jmps.180420 However, observations of the Fe-enrichment of olivine T. Nozaka, nozaka@ cc.okayama-u.ac.jp Corresponding author 220 T. Nozaka SWJapan RM TM AS OS 100 km Ultramafic complex SM Suo Belt (230-160 Ma) SK Sangun high-P/T OE HP Renge Belt (330-280 Ma) metamorphic belt Figure 1. Distribution of ultramafic complexes and the Sangun high-P/T metamorphic belt (Renge and Suo Belts) in SW Japan (Ishiwatari, 1989, 1990; Nishimura, 1998; Takeuchi, 2002; Isozaki et al., 2010). Abbreviations for ultramafic complexes: AS, Ashidachi; HP, Happo; OE, Oeyama; OS, Ohsayama; RM, Ryumon; SK, Sekinomiya; TM, Tari-Misaka; WS, Wakasa. Color version is available online from https://doi.org/10.2465/jmps.180420. except for some examples (Murata et al., 2009a, 2009b; Nozaka, 2003, 2005; Khedr and Arai, 2012). The tempo- ral sequence of these geologic events was summarized by intense low-T serpentinization or overprinting metamor- Nozaka (2014a). phism. In addition to reporting new occurrence of the retro- ANALYTICAL PROCEDURES grade ferroan olivine in the Oeyama ophiolite, SW Japan, I show here evidence for the volume diffusion of ele- Polished thin sections of serpentinized peridotites from ments in olivine grains under the conditions of high-T the Oeyama ophiolite were prepared for microscopic ob- serpentinization. This compositional modification of oliv- servations and chemical analyses of minerals. Several ine could be a useful indicator of tectonothermal history samples from the Ryumon ultramafic body in the Sanba- of serpentinized forearc peridotites. gawa metamorphic belt exposed in the Kii Peninsula were analyzed for comparison. Quantitative microprobe GEOLOGICAL SETTING analyses were carried out using an electron probe micro- analyzer, JEOL JXA-8230 at Okayama University with The Oeyama ophiolite is a collective name of ultramafic an accelerating voltage of 15 kV and a probe current of complexes exposed in the Renge high-P/T metamorphic s s o s o belt of SW Japan (Fig. 1; Ishiwatari, 1989, 1990; Nishi- Fe, and Mg), 60 s for minor elements (Mn, Ni, and Ca) of mura, 1998; Isozaki et al., 2010). In the present article, olivine, and 10 s for the all element of serpentine; back- the Happo complex is included in the Oeyama ophiolite ground counting time was 5 s for all the cases. Standards despite its long distance from the other complexes, be- used were natural or synthetic oxides and silicates. The cause they are similar in geological and geochemical applied matrix corrections followed the procedures of characteristics (Takeuchi, 2002; Nozaka, 2005; Khedr Bence and Albee (1968), using alpha factors of Naka- and Arai, 2010; Nozaka, 2014a). mura and Kushiro (1970). Identifcation of serpentine The peridotites of the Oeyama ophiolite are consid- minerals of representative samples was confirmed with ered to be residual mantle peridotites and cumulates from Raman shift spectra, using a micro-Raman spectrometer, basaltic melt originated at a Paleozoic supra-subduction JASCO NRS-3100 at Okayama University, with 488 nm zone or forearc mantle wedge (Ishiwatari, 1989; Arai and laser excitation, a 100 × microscope objective lens, and a Yurimoto, 1995; Ishiwatari and Tsujimori, 2003; Khedr diffraction grating with 1800 grooves/mm. and Arai, 2010; Nozaka, 2014a). They were subjected to two stages of mylonitization, which could be associated RESULTS with metasomatic hydration at 700-800 °C and later high-T serpentinization at 400-600 °C (Nozaka, 2005; As I have previously reported the petrographic observa- Nozaka and Ito, 2011; Nozaka, 2014a). Many of the ser- tions of the serpentinized peridotites from the Oeyama pentinized peridotite bodies of the Oeyama ophiolite ophiolite (Nozaka, 2005; Nozaka and Ito, 2011; Nozaka, were intruded by Cretaceous or Paleogene granitic rocks 2014a), I just briefly describe here an outline of the char- and consequently thermally metamorphosed (Arai, 1975; acteristics of olivine and serpentine. Primary olivine of Compositional variation of olivine in high-temperature serpentinites 221 the peridotites shows modes of occurrence with variable porphyroclasts with decreasing distance from antigorite effects of plastic deformation: relatively coarse, equant in serpentinites from the Happo complex. Such a compo- grains with subgrain boundaries; ^cleavable olivine′ with sitional variation is unclear in other ultramafic complexes prominent parting planes; and elongated elliptic porphyr probably because of intense low-T serpentinization and/ oclasts distributed in a matrix of fine-grained recrystal. or contact metamorphism. The plots of olivine composi- lized olivine and lepidoblastic antigorite. The primary ol- tion were obtained avoiding the bright halos and streaks ivine grains are serpentinized more or less from rims and in back-scattered electron images mentioned above but along parting planes or fractures. Antigorite blades pen- could be affected by modification along invisible frac- etrate the porphyroclastic or cleavable olivine and are as- tures or subgrain boundaries to some extent as discussed sociated with chlorite, diopside and/or brucite in the lep- later. Nickel contents show no systematic variation with idoblastic matrix of mylonites. These antigorite blades distance unlike Fo and Mn (Fig. 3). have been interpreted to be a product of high-T serpenti- In the diagrams of NiO versus Fo and MnO versus nization (Nozaka, 2005). Lizardite, which is a product of Fo contents, a cluster of plots from each sample exhibits low-T serpentinization, statically replaces olivine to form a trend of compositional variation from the magnesian mesh texture with tiny grains of opaque minerals (mostly primary olivine to the most ferroan retrograde olivine magnetite) and submicroscopic brucite. The mesh-form- (Fig. 4). The NiO-Fo plots are similar, though more scat- ing lizardite is replaced by metamorphic olivine in contact tered, to the ‘mantle olivine array’ defined by Takahashi aureoles around granitic intrusions. The contact metamor- et al. (1987). The trend is much evident in MnO-Fo dia- phic olivine occurs as discrete small grains or epitaxially gram and designated here as ‘retrograde trend' (Fig. 4b). grown rims over primary olivine grains, coexisting with Prograde olivine occurs in the contact aureoles of antigorite at a low grade zone (Zone I) and with talc, most of the ultramafic complexes of the Oeyama ophio- anthophyllite and orthopyroxene at higher grade zones lite. In a low-grade zone (Zone II), the prograde olivine (Nozaka, 2003, 2005, 2011; Nozaka and Ito, 2011; No- commonly has highly magnesian compositions (Arai, zaka, 2014b). 1975; Nozaka, 2005; Nozaka and Ito, 2011; Khedr and Reflecting the metamorphic history of the Oeyama si s s n n ophiolite, at least two generations of antigorite can be prograde olivine with relatively ferroan compositions observed: i.e., retrograde antigorite formed by high-T ser- (Nozaka, 2003). The ferroan prograde olivine is similar pentinization and prograde antigorite formed by meta- in Fo content to the retrograde olivine, but is different in morphism after low-T serpentinization. Similarly, I will highly variable contents of NiO and MnO (Fig. 4; No- refer to the olivine coexisting with the former antigorite zaka, 2003, 2010). Relatively ferroan olivine formed by or affected by the high-T serpentinization as ‘retrograde regional metamorphism has been reported from the San- olivine’ and that coexisting with the latter antigorite as bagawa metamorphic belt, and it is also different in NiO 'prograde olivine'. and MnO contents from the retrograde olivine (Kunugiza, The retrograde olivine is common in serpentinite 1980, 1982; Nozaka, unpublished data). mylonites from the Happo and Wakasa complexes and Retrograde antigorite coexisting with retrograde ol- locally found outside the contact aureole of the Ohsa- ivine occurs as elongated blades surrounding or penetrat- yama complex. This olivine does not form discrete grains ing primary olivine (Figs. 2e-2i). The antigorite blades but occurs as a part of chemically modifed primary oliv. commonly lack magnetite inclusions unlike lizardite ine in contact with retrograde antigorite blades. It cannot (Figs. 2d, 2g, and 2h) and locally has the intercalation be distinguished from primary olivine under the optical of elongated magnetite, chlorite or brucite. A small microscope but is visible by a high electron back-scatter- amount of awaruite in contact with the antigorite blades ing intensity resulting from an Fe-rich composition (Figs. was observed in one sample (HP21). Retrograde antigor- 2a-2f, 3, and Table 1). Such an enrichment of Fe is lack- ite has aluminous compositions indicative of the incorpo- ing in olivine at any portions in contact with lizardite ration of chlorite components, and lacks evidence for mix- (Figs. 2c and 2f). Back-scattered electron images show ing with brucite unlike the case of lizardite (Figs. 5a and bright halos around some small antigorite blades that are 5c), which shows a brucite-mixing trend as reported from seemingly inclusions in olivine but possibly fracture-fill- many localities (e.g., Beard et al., 2009). However, the ings cut by the thin-section surface (Fig. 2c) and locally Al-enrichment is less significant in antigorite apart from show a small amount of bright streaks ~ 1 μm wide (not matrix chlorite as the case of fracture-filling antigorite in shown) within olivine grains. olivine porphyroclasts (Fig. 2d). Lizardite is commonly Figure 3 shows a systematic decrease of Fo [= 100 x less aluminous than retrograde antigorite, excepting bas- Mg/(Mg + Fe)] and increase of Mn contents in olivine tite-forming Al-rich lizardite (not shown in the figures). 222 T. Nozaka Figure 2. Mode of occurrence of olivine and antigorite. Abbreviatio ite; Ol, olivine; P-Ol, prograde olivine. (a) Photomicrograph of an olivine erpentinite mylonite from Happo (crossed porphy polars, sample# HP40). (b) Back-scattered electron image of the shown in (a). (c) Back scatteredelectron dynamically recrystallized olivine neoblasts and lepidoblastic antigorite of & rpentinite mylonite from Happo (crossed polars, sample# HP11). (f) Back-scattered electron image of a part of (e). Color version is available online from https://doi.org/10.2465/jmps.180420. Prograde antigorite occurs in contact metamorphic ite is characteristically smaller in size and in aspect ratio zones where prograde olivine occurs. This type of antigor- (length/width) than retrograde antigorite, and typically 223 O OI b1t 1.0mm Figure 2 (Continued). Mode of occurrence of olivine and antigorite. Abbreviations for minerals: Atg, antigorite; Chl, chlorite; Liz, lizardite; Mgt, magnetite; Ol, olivine; P-Ol, prograde olivine. (g) Photomicrograph of retrograde antigorite and lizardite replacing olivine in a gently deformed serpentinite from Wakasa (plane-polarized light, sample# WS0301). (h) Back-scatered electron image of a part of (g). (i) Photo- micrograph of écleavable olivine’ surrounded and penetrated by antigorite from a low-grade contact aureole (Zone I) in the Wakasa e d o r ( ( go pd e o s o psa ( (o #s sd pss xs from a low-grade contact aureole (Zone II) in the Oeyama complex (crossed polars, sample# OE0817). (l) Back-scattered electron image of a part of (k). Color version is available online from https://doi.org/10.2465/jmps.180420. 224 T. Nozaka 90.0 91.0- 89.0 I1 (a) HP40 88.0 (b) HP11 87.0 品 1500 I1c dd I1α Figure 3. Variation of forsterite (Fo), M M Mn, and Ni contents of olivine 1000- with the distance from the contact 1000- with (a) fracture-flling antigorite (sample #HP40 shown in Fig. 2b), and (b) antigorite matrix (sam- ple #HP 11 shown in Fig. 2f). 3500 Curves indicate possible diffusion profiles drawn to fit a set of plots 3500 (circle) with high Fo and low Mn values at each distance. Other plots (square), which show a considera- ble deviation from the diffusion profile, probably represent modifi- 口口口 1 cation along invisible fractures or 2500- subgrain boundaries. Color version 100 200 300 400 100 200 300 400 is available online from https:// Distance from antigorite vein (μm) Distance from antigorite matrix (μm) doi.org/10.2465/jmps.180420. forms fine-grained aggregate with subordinate amounts of grade antigorite in the peridotites of the Oeyama ophio- magnetite (Figs. 2k and 2l). Antigorite of regional meta- lite. Further evidence is that plots of Xre of the most morphic origin from the Ryumon ultramafic body in the ferroan olivine-antigorite pair in each sample are consis- Sanbagawa belt is also smaller in grain size than the ret- tent with those of coexisting pairs from other localities rograde antigorite of the Oeyama ophiolite and shows (Fig. 6). The enrichment of Fe in olivine porphyroclasts granoblastic texture with metamorphic olivine (Kunugiza, and the deficiency of magnetite associated with retro- 1980, 1982), suggesting their contemporaneous forma- grade antigorite indicate a reducing condition and support tion. The prograde antigorite, including that of the Ryu- the hypothesis that Mg-Fe interdiffusion in olivine is ef- mon body, typically has more magnesian compositions fective under the conditions of high-T serpentinization (XMg > 0.96) than the retrograde antigorite (XMg = 0.94- proposed by Evans (2010). Although a small amount of 0.96) (Fig. 5d). The antigorite blades that have a typical elongated magnetite is intercalated in some antigorite shape, size, and mode of occurrence of retrograde antigor- blades, they are not necessarily formed contemporane- ite occur in low-grade contact aureoles as well (Figs. 2i ously. Because brucite showing a similar mode of occur- and 2j) and have compositions scattering in both the fields rence to magnetite was observed in some serpentinite of typical retrograde and prograde antigorite (Fig. 5d). mylonites, the possibility that the magnetite replaced brucite, as suggested by Bach et al. (2006) and Beard DISCUSSION et al. (2009), cannot be ruled out. The variations of Fo content with distance from an- Compositional modification of olivine tigorite (Fig. 3) suggest that Mg-Fe interdiffusion took place within the olivine porphyroclasts, i.e., between The textural relationship strongly suggests the contempo- the cores keeping primary compositions and the rims of raneous formation of the retrograde olivine and retro- retrograde olivine coexisting with the retrograde antigor- Compositional variation of olivine in high-temperature serpentinites 225 Table 1. Representative microprobe analyses of olivine and serpentine Sample HP40 HP40 HP40 HP40 HP11 HP11 HP11 WS0615 WS0615 WS0615 WS0615 WS0615 OE0817 Mineral 10 10 Atg Liz 10 10 Atg 10 10 10 Atg Liz Atg Type Pri R R Pri R R-P Pri R P R-P P SiO2 40.51 40.60 42.66 42.36 40.06 39.89 43.02 40.21 40.06 40.82 43.24 39.76 44.94 TiO2 pu nd 0.02 0.03 pu nd 0.01 pu pu nd 0.01 0.00 0.00 nd Al2O3 pu 2.22 0.26 pu pu 2.54 nd nd nd 1.16 0.01 0.08 Cr2O3 nd nd 0.20 0.00 nd nd 0.24 nd pu pu 0.33 0.00 0.00 FeO* 10.22 13.06 3.90 2.10 9.79 11.62 3.59 9.50 10.48 5.81 2.49 1.96 1.53 MnO 0.15 0.25 0.05 0.10 0.14 0.22 0.04 0.13 0.18 0.04 0.05 0.05 0.00 NiO 0.39 0.35 0.22 0.07 0.38 0.37 0.16 0.39 0.40 0.22 0.06 0.25 0.05 MgO 49.16 46.33 36.65 38.80 50.03 47.63 36.76 49.98 48.69 53.12 37.61 39.25 39.57 CaO 0.00 0.01 0.00 0.11 0.00 0.00 0.01 0.00 0.01 0.00 0.01 0.04 0.00 Na2O pu nd 0.00 0.01 nd nd 0.04 pu pu pu 0.00 0.01 0.01 K20 nd nd 0.01 0.02 pu nd 0.00 nd nd nd 0.00 0.01 0.00 Total 100.45 100.59 85.94 83.86 100.40 99.73 86.41 100.21 99.81 100.00 84.96 81.33 86.18 0= 4.000 4.000 7.000 7.000 4.000 4.000 7.000 4.000 4.000 4.000 7.000 7.000 7.000 Si 0.992 1.004 2.023 2.043 0.981 0.991 2.023 0.985 0.989 0.984 2.056 1.988 2.093 Ti pu pu 0.001 0.001 nd nd 0.000 pu pu pu 0.000 0.000 0.000 Al nd nd 0.124 0.015 nd nd 0.141 nd nd nd 0.065 0.000 0.004 Cr nd nd 0.007 0.000 nd pu 0.009 nd nd nd 0.013 0.000 0.000 Fe 0.209 0.270 0.154 0.085 0.201 0.241 0.141 0.195 0.217 0.117 0.099 0.082 0.060 Mn 0.003 0.005 0.002 0.004 0.003 0.005 0.002 0.003 0.004 0.001 0.002 0.002 0.000 Ni 0.008 0.007 0.009 0.003 0.007 0.007 0.006 0.008 0.008 0.004 0.002 0.010 0.002 Mg 1.795 1.709 2.591 2.790 1.827 1.764 2.578 1.825 1.793 1.909 2.666 2.926 2.747 Ca 0.000 0.000 0.000 0.006 0.000 0.000 0.001 0.000 0.000 0.000 0.001 0.002 0.000 Na nd nd 0.000 0.001 nd nd 0.004 nd nd pu 0.000 0.000 0.001 K pu nd 0.001 0.001 pu pu 0.000 pu nd pu 0.000 0.000 0.000 Total 3.008 2.996 4.911 4.949 3.019 3.009 4.904 3.015 3.011 3.016 4.905 5.012 4.906 XMg 0.896 0.863 0.944 0.971 0.901 0.880 0.948 0.904 0.892 0.942 0.964 0.973 0.979 * Total iron as FeO; nd, not determined; XMg = Mg/Mg + Fe) molar ratio. Mineral abbreviations: Atg, antigorite; Liz, lizardite; Ol, olivine. Wakasa (WS) and Oeyama (OE). ite. Relatively low Fo values distributed below the diffu- ism. Since volume diffusion is a process more sluggish sion profile curves (Fig. 3) could be the effects of hidden than grain/subgrain boundary diffusion or pipe diffusion, (not appeared on the thin-section surface) irregular small the serpentinization of the Oeyama ophiolite could pro- fractures filled with antigorite as shown in Figure 2c. ceed at a higher-T or over a longer period than the ser- Stripes of Fe-enriched olivine have been reported and pentinization accompanied by the ferroan olivine stripes. interpreted as a result of Mg-Fe interdiffusion along sub- Because Mg-Fe volume diffusion in olivine looks unre- grain boundaries, or pipe diffusion along dislocations, alistic below 400 °C as discussed later, it probably took within deformed olivine crystals (Kitamura et al., 1986; place under the conditions of high-T serpentinization. Ando et al., 2001; Murata et al., 2009a, 2009b; Plumper Figure 3 indicates that Mn is enriched in retrograde et al., 2012). Although the bright streaks visible in back- olivine as well as Fe. As pointed out by Evans (2010) and scattered electron images suggest that an Fe-enrichment Plimper et al. (2012), this enrichment can be explained along linear defects within olivine crystals took place in by the average distribution coeffcient of Mn established the Oeyama ophiolite as well, their amount and frequen- by Trommsdorff and Evans (1974). For example, the XMn cy are just minor and volume diffusion, which forms a [= Mn/(Mg + Fe + Mn + Ni)] of the most ferroan olivine systematic compositional variation depending on the dis- in two samples (HP11 and HP40) are 0.0019-0.0026, and tance from antigorite, is probably the dominant mechan- applying the Mn distribution coefficient Kp (antigorite/ 226 T. Nozaka 0.6 0.3- (a) b 0.5- 0.4 MOA 0.2 % MOA N 0.3 O!N MnO 壬1o 0.2 PFO 0.1- ●HP40 0.1 HP11 WK0615 壬1α 0.0 0.0- 86 87 88 89 90 91 86 87 88 89 90 91 Fo mol% Fo mol% Figure 4. (a) NiO versus Fo contents and (b) MnO versus Fo contents of olivine in contact with antigorite in serpentinite mylonites (sample# by high-T serpentinization and intracrystalline element diffusion. Compositional fields of olivine from the literature are shown for compar- ison: MOA, mantle olivine array (Takahashi et al., 1987); PFO, prograde metamorphic ferroan olivine from Zone II of contact aureoles, SW Japan (Nozaka, 2003, 2010). Iron-rich prograde metamorphic olivine from the Ryumon peridotites plot within PFO or significantly out of the retrograde trends, or outside the scale ranges (Kunugiza, 1980, 1982; Nozaka, unpublished data). Color version is available online from https:/doi.org/10.2465/jmps.180420. olivine) = 0.18 (Trommsdorff and Evans, 1974) to these ite, as pointed out by Plumper et al. (2012). In fact, the olivine, the calculated Xmn of antigorite are 0.0003- coexistence of awaruite with antigorite, though a small 0.0005. These values are consistent with the XMn of an- amount, was observed in at least one sample from the alyzed antigorite: 0.0002-0.0008 with an average of Oeyama ophiolite. 0.0005 (24 analyses). It is in no doubt that the successive zonal arrange- The Ni contents of olivine porphyroclasts do not ments of olivine compositions in NiO-Fo and MnO- show such a systematic variation as Fo and Mn (Fig. 3). Fo diagrams (Fig. 4) are the results of intracrystalline In the same manner as Mn, using the Kp = 0.65 (Tromms- element diffusion between the most ferroan retrograde dorff and Evans, 1974), XNi was calculated for antigorite olivine and primary olivine. It is the case as well for sam- coexisting ferroan olivine to be 0.0021-0.0027. These val- ples from a low-grade zone (Zone I) of contact aureole ues are not inconsistent with the analyzed values of anti- (e.g., HP11 and WS0615 shown in Fig. 4), in which high- gorite: 0.0015-0.0034, and therefore retrograde olivine ly magnesian prograde olivine also occurs (Fig. 2j), sug- and retrograde antigorite probably tend to keep their equi- gesting the existence of relict retrograde olivine in the librium for Ni as well. However, such a diffusion profile contact aureole. Thermal metamorphism with tempera- as the case of Mn cannot be defined for Ni in the olivine ture conditions below 480 °C, which were estimated for porphyroclasts (Fig. 3). Olivine, which has enrichments of that zone (Nozaka, 2011), looks to have no effect on the Fe and Mn showing deviation from the diffusion profiles compositions of retrograde olivine. at a significant distance from antigorite probably due to We can recognize the effects of high-T serpentiniza- fluid infiltration along fractures or dislocations, shows an tion on the compositional modification of olivine from increase and decrease of Ni from the primary composition textural evidence as shown in some samples (Figs. 2b, (e.g., the three data on the extreme right of Fig. 3a). This 2f, and 3) and as previously reported (Murata et al., opposite behavior of Ni suggests a variation of Ni content 2009a, 2009b; Plimper et al., 2012). However, such a of fluid during high-T serpentinization. The increase of case is uncommon because of the overprinting effects Ni probably results from its distribution between olivine of intense low-T serpentinization and later metamor- and antigorite (Kp = 0.65 according to Trommsdorff phism. The occurrence of ferroan olivine is an indicator and Evans, 1974), whereas the decrease of Ni could be of high-T serpentinization, but we cannot exclude other explained by the formation of Ni-rich phase, e.g., awaru- possibilities for the formation of ferroan olivine in ser- Compositional variation of olivine in high-temperature serpentinites 227 pentinized peridotites of ophiolites: i.e., crystallization (Figs. 2i and 2j). differentiation of basaltic magma, deformation-induced fGakken School,KonoBuilding,27-17Takashima-cho,Numazu410-0056,aan pipe diffusion, metasomatic fuid infiltration, and pro- equilibration compositions are maintained. Retrograde an- grade metamorphism. In these cases, a set of NiO-Fo tigorite is relatively rich in Fe resulting in low Xmg (Figs. and MnO-Fo diagrams (Fig. 4) could be useful to distin- 5b and 5d). The contrastingly high XMg of prograde anti- guish the ferroan olivine of different origin. In a given gorite is likely inherited from lizardite, a reactant from sample from the mantle section of ophiolites, composi- which the prograde antigorite has formed. During low-T tions of olivine produced or affected by high-T serpenti- serpentinization, Fe in olivine is mainly incorporated into nization are expected to be plotted around the retrograde magnetite (e.g., Nozaka, 2003; Evans, 2010), some of trends of Fo-MnO without no significant deviation of Minoo-higashi High School, 5-4-63 Ao-Gein, Minoo 552-0025,Japan NiO from the mantle olivine array (Fig. 4). From this DepartmentfGesciencehmaneiversityatsu90504,an type of olivine, ferroan olivine produced by other pro- nesian compositions of lizardite. Although the XMg of cesses could be different in some points: a progressive lizardite and antigorite varies depending on primary oliv- depletion of NiO and a gentle enrichment of MnO with ine or bulk rock compositions, the retrograde antigorite decreasing Fo contents by magmatic differentiation directly formed after primary olivine should be richer in (Takahashi et al., 1987); no significant change of NiO Fe than prograde antigorite after lizardite in any samples. and MnO contents by high-T pipe diffusion in anhydrous, However, some relict blades of retrograde antigorite in not serpentinized peridotite (Ando et al., 2001); much contact aureoles have a slightly magnesian compositions less extent of Fo, NiO, and MnO variations by amphib- probably resulting from a reequilibration process during olite-facies metasomatism (Nozaka, 2014a); and highly E-mail address: imaoka@yamaguchi-u.ac.jp (T. Imaoka). variable NiO and MnO contents with a large deviation ture of retrograde antigorite of the Oeyama ophiolite is from the retrograde trend and the mantle olivine array an Al-enrichment (Figs. 5c and 5d). This could be an in- by prograde contact/regional metamorphism (Kunugiza, dication of increasing chlorite component in antigorite 1980, 1982; Nozaka, 2003, 2010). under high-T conditions (Padron-Navarta et al., 2013). Before discussion about T conditions, however, a caution Retrograde vs. prograde antigorite is needed for whether or not the antigorite coexists with chlorite; in fact, antigorite in olivine fractures apart from Retrograde antigorite would be easily found if it shows matrix chlorite, tends to be deficient in Al. textural and chemical equilibration with retrograde oliv- * Corresponding author. Tel.: +81 83 933 5765; fax: +81 83 933 5273. ine, but retrograde olivine grains might disappear because ite requires the diagnostic characteristics: a shape of elon- of intense low-T serpentinization. The question is wheth- gated blade, the absence of a significant amount of mag- er it is possible to discriminate between retrograde and netite inclusion, and a composition richer in Fe than that prograde antigorite in the latter case. of lizardite. These characteristics would provide empiri- The two types of antigorite have differences in some cal evidence for high-T serpentinization even in intensely microscopic features. Retrograde antigorite commonly serpentinized peridotites from ophiolites. However, the occurs as relatively large blades with high aspect ratios possibility of chemical disturbance in prograde metamor- and without magnetite inclusions (Figs. 2d, 2g, and 2h). phic rocks should be carefully investigated. - s o gates with minor amounts of magnetite in contact aur- Tectonic implication eoles (Figs. 2k and 2l) or occurs as relatively small blades forming an equilibration texture; e.g., granoblastic tex- The diffusion profiles of olivine porphyroclasts (Fig. 3) ture with prograde metamorphic olivine (e.g., Kunugiza, provide a constraint for tectonic model of the forearc 1980). However, retrograde antigorite also occurs as a mantle wedge (Fig. 7; Schmidt and Poli, 1998; Peacock small blade (Fig. 2d), and the shape or size may not be and Wang, 1999; Stern, 2002; Hacker et al., 2003; Rey- concluding evidence for identification. nard et al., 2011; Nozaka, 2014a). The illustrated diffu- 0024-4937/S - see front matter @ 2013 Elsevier B.V. All rights reserved. sion profiles differ between the two samples (Fig. 3) in a given by metamorphic zonal mapping. In the contact au- reflection of primary olivine composition, and possibly in reoles of SW Japan, prograde antigorite occurs in Zone addition, crystallographic orientation, degree of crystal II, which is defined by the first appearance of prograde deformation and T condition. Nevertheless, similar diffu- metamorphic olivine (e.g., Nozaka, 2003, 2011). How- sion distances are determined: ~ 35 μm for sample HP40 ever, this is not a condition for excluding the occurrence and ~ 45 μm for sample HP11, which allow us to esti- of retrograde antigorite, because it also occurs in Zone II mate the time necessary for the diffusion. 228 T. Nozaka Brc 0.4 been intensely metasomatized (Ishimaru et al., 2009). These 105 Ma (a) Figure 5. Serpentine and chlorite OPAtg compositions calculated on the ba- PAtg (RM) R-P Atg sis of 7 oxygen atoms: (a) Mg + 5 Liz ±Brc Fe* versus Si (Fe*, total iron as Fig. 1. Index map of Southwest Japan (A), geotectonic divisions of post-Triassic accretionary complexes in the western part of SW Japan (B), and distribution of Jurassic and Cretaceous Fe2+); (b) Fe* versus Si; (c) Al ver- II = 7) sus Si; (d) Al versus Mg/(Mg + Fe* O 0.2- Brc trend - Fe*), in which retrograde- and pro- Atg + Cina Tsuchiya et al., 2005). Studies of these hornblende peridotite xenoliths grade-antigorite composition fields rocks can contribute significantly to our understanding of the chemical in serpentinites from SW Japan are 0.1- highlighted. All the plotted serpen- Chl trend tine and chlorite are replacing oliv- ine, coexisting with olivine, or 0.0- forming mylonite matrix, and bas- tite-forming minerals are not plot- Si (0 = 7) Si (0= 7) ted. Arrows indicate mixing trend Atcin with brucite or chlorite. Abbrevia- 1.0- (c) (d) tions: Brc, brucite; Chl, chlorite; 0.2 Cln, clinochlore; Liz, lizardite; and physical processes operating in the subcontinental lithosphere be- 0.8- Retrograde Ol, olivine; P Atg, prograde anti- -antigorite gorite; R Atg, retrograde antigor- ite; R-P Atg, retrograde antigorite = a :() i s o = o e 's = on = m 's sa = ) s () 7 Prograde that has a typical shape, size and = 0) 1995, 2000), as well as the high-Mg, Nb-rich lamprophyres that are antigorite mode of occurrence in prograde 0.4- A contact aureoles. Prograde antigor- ite from the regionally metamor- phosed Ryumon (RM) peridotite 0.2- body (Kunugiza, 1982; Nozaka, unpublished data) plots for com- [Brc Brctrend plement. Color version is avail- 0.0+ 白 2Atg 0.92 0.94 0.96 0.98 1.00 able online from https://doi.org/ Si (O = 7) Mg/(Mg + Fe*) 10.2465/jmps.180420. Arc crust High-T serpentinization 600°℃ Exhumation 800°℃- Mantle wedge characteristics that include Si02 ≥ 56 wt.%, Al203 ≥ 15 wt.%, Mg0 1000℃ interpreted as indicating both the presence of residual garnet and Amphibolite-facies metasomatism. 100°0 0.00 0.10 0.20 Fe*(Mg + Fe*) oline Figure 6. Plots of Fe*/(Mg + Fe*) of the most ferroan olivine- antigorite pair in each sample (stars) from the Oeyama ophiolite Fluidinfiltration (sample# HP11, HP40, HP44, WS0615, OS0502). The field of olivine-antigorite pairs compiled by Evans (2010) is also shown Figure 7. Schematic model of the sequence of metasomatism, hy- for comparison. Color version is available online from https:// dration and exhumation of the forearc mantle peridotites repre- doi.org/10.2465/jmps.180420. sented by the Oeyama ophiolite based on a synthesis of some previous works (Schmidt and Poli, 1998; Peacock and Wang, The diffusion time was calculated on the basis of the 1999; Stern, 2002; Hacker et al., 2003; Reynard et al., 2011; equation: x ~ √Dt, where x, D and t is diffusion distance, Nozaka, 2014a). See Nozaka (2014a) for discussion. Color version is available online from https://doi.org/10.2465/jmps. diffusion coefficient and time, respectively, using a diffu- igneous rocks in and around the Korean Peninsula and the western part of SW Japan (C). Fig. 1c is compiled from Kee et al. (2010), Teraoka and Okumura (2003), and Wu et al. sion distance of 40 μm (Fig. 3) and the diffusion coeff- cient based on the global equation of Dohmen and Chak- raborty (2007a, 2007b). The diffusion coefficient for the equation, and it is necessary to subtract log 6, i.e., direction of [Oo1] of olivine is obtained by the global ~ 0.78 order of magnitude for [100] and [010] according 229 to Dohmen and Chakraborty (2007a, 2007b). However, be useful to distinguish retrograde or retrogressively the difference of D below 700 °C is less than 0.3 order of modified olivine from ferroan olivine of different origin, magnitude (Dohmen et al., 2007). Actually, the difference even if intense low-T serpentinization has obscured the is considered to be much less because the interference early record of the diffusion process. Retrograde antigor- colors under the microscope indicate that the directions and metabasites, and radiometric ages indicate metamorphism in dikes are dominated by hornblende and plagioclase with <2 vol.% of elongated blade, lacking a significant amount of mag- tain the minimum duration of cooling during or after netite inclusion, and having a composition richer in Fe high-T serpentinization, T conditions were supposed to than that of lizardite. These characteristics are expected be close to the maximum estimation for high-T serpenti- to provide empirical evidence for high-T serpentinization nization of the Oeyama ophiolite (Nozaka, 2014a). The even in intensely serpentinized peridotites from ophio- rounding granitoids of the San-yo and Ryoke belts. Hornblende and lites. Approximate estimation of time required for the ob- ~ 4.0 × 107 years at 500 °C and 0.5 GPa. An unreason- served diffusion profiles of olivine porphyroclasts sug- ably long time (> 4.5 × 10? years) was estimated at 400 gests a long residence time of the forearc peridotites °C, suggesting the observed intracrystalline element dif- within the serpentinized mantle wedge following rapid topes of the Early Cretaceous igneous rocks from the Kinki district of exhumation immediately after the amphibolite-facies of high-T serpentinization. metasomatism. In contrast, much shorter diffusion time of tens to hundreds of years has been estimated for the diffusion ACKNOWLEDGMENTS profile caused by the fuid injection that formed oliv- ine-phlogopite veins in the Oeyama ophiolite and inter- I thank Dr. Masaki Mifune for permission to use a Raman preted as a result of rapid exhumation after amphibolite- spectrometer in his laboratory at Okayama University. facies metasomatism of the forearc mantle wedge (No- Comments from reviewers, J. Ando and E. Takazawa zaka, 2014a). That short time was obtained using D of have improved the manuscript. This study was finan- Jaoul et al. (1995), which is one or two orders of magni- cially supported by JSPS KAKENHI Grant Number tude larger than that of Dohmen and Chakraborty (2007a, JP16K05611. 2007b). Using the latter D in this study, the time was recalculated to be ~ 5.0 × 10² to 8.0 × 103 years at 750- SUPPLEMENTARYMATERIALS 800 °C, which is still an extremely short period in the geologic time frame. Although exact time required for Color versions of Figures 1-7 are available online from fusions. Errors and the analytical precision of the ICP-MS analyses are https://doi.0rg/10.2465/jmps.180420. because of the uncertainties of D, T, and pressure condi- tions, the approximate estimation suggests that the cool- REFERENCES ing duration under the conditions of high-T serpentiniza- tion was at least three to four orders of magnitude longer Ando, J., Shibata, Y., Okajima, Y., Kanagawa, K., Furusho, M. and Tomioka, N. 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Journal of Mineralogical and Petrological Sciences, Volume 115, page 431-439, 2020 A new occurrence of okhotskite in the Kurosegawa belt, Kyushu, Japan: the okhotskite + Mn-lawsonite assemblage as a potential high-pressure indicator Wataru YABUTA and Takao HIRAJIMA Department of Geology and Mineralogy, Division of Earth and Planetary Sciences, Graduate School of Science, Kyoto University, Kyoto 606-8502, Japan We present the first report of okhotskite in a lawsonite-blueschist-subfacies metachert of the Hakoishi subunit, Kurosegawa Belt, Kyushu, Japan, which was metamorphosed at peak temperatures and pressures of 200-300 °C sures than that of previously documented okhotskite with available pressure estimations. Textural relationships indicate that okhotskite formed during peak metamorphism in equilibrium with piemontite, Na pyroxene, mag- nesioriebeckite, braunite, and hematite. Okhotskite shows a significant variation in Fe:Mn ratio (Feto/Mntot = 0.13-0.56) and a following average empirical formula; (Ca7.62Mn2.16)≥7.78(Mn2.71Mg1.29)≥4.00(Mn4.13Fe2.26Al1.36 V323 Tio.02)≥8.0oSi11.86O44.02(OH)16.98. Raman spectra of okhotskite are reported for the first time and show char- acteristic peaks at 362, 480, and 563 cm-1. The stability relationships between okhotskite and other Mn-bearing minerals, such as piemontite, sursassite, spessartine, braunite, and Mn-bearing lawsonite, are examined using a revised Schreinemakers? analysis. The obtained petrogenetic grid provides tight constraints on the P-T relation- ship of natural mineral assemblages observed in Mn-bearing cherts within epidote-blueschist-grade and law- sonite-blueschist-grade. Furthermore, this petrogenetic grid predicts that the assemblage of okhotskite and Mn- bearing lawsonite should be stable at higher pressures. The higher-pressure stability suggests that highly oxi- dized Mn-bearing metacherts can transport water and buffer oxygen in the deeper parts of subduction zones, given that okhotskite and Mn-bearing lawsonite contain high water contents (6.9 and 11.3 wt% H2O, respec- tively) and trivalent manganese. Keywords: Okhotskite, Lawsonite-blueschist subfacies, Metachert, Kurosegawa Belt, Schreinemakers’ analy- sis, Raman spectra INTRODUCTION do, Japan (Togari and Akasaka, 1987). Okhotskite has also been reported from metamor- Pumpellyite group of minerals have a general chemical phosed manganese deposits in the Sanbagawa (Mina- formula of [W&X4YsZ12O56-n(OH)n] (Passaglia and Got- kawa, 1992) and Chichibu (Minakawa et al., 2008) belts tardi, 1973) and are common hydrous minerals in low- in Shikoku, Japan; metamorphosed manganese oxide grade metamorphic rocks. Root-names of the group have ores of the Precambrian Sausar Group, India (Dasgupta been defined based on the dominant cation in the Y-site: et al., 1991); and manganese deposits in the Askiz ore pumpellyite with A13+ (Palache and Vassar, 1925), julgol- district, Khakassia, Russia (Kassandrov and Mazurov, dite with Fe3+ (Moore, 1971), shuiskite with Cr3+ (Ivanov 2009). These reports include the coexistence okhotskite et al., 1981), okhotskite with Mn3+ (Togari and Akasaka with braunite (Mn2+Mn3+SiO12) and/or bixbyite (Mn2O3), 1987), and poppite with V3+ (Brigatti et al., 2006). and the presence of trivalent Mn suggests that the mineral Okhotskite, the Mn3+-dominant member, was first report- records a strongly oxidizing environment. Furthermore, ed from the Kokuriki mine, in the Tokoro Belt, Hokkai- okhotskite is expected to provide a considerable part of the water budget of impure metacherts in the subduction doi:10.2465/jmps.190831 W. Yabuta: yabuta.wataru.25n@kyoto-u.jp Corresponding author zone, because of its high H2O content (6.89 wt% H2O; T. Hirajima: hirajima.takao.4u @ kyoto-u.ac.jp Togari and Akasaka, 1987).
432 W. Yabuta and T. Hirajima In considering phase relations among such highly (a) sw Japan oxidized Mn minerals, oxygen fugacity (fo2) has been considered as a critical factor in addition to pressure (P) and temperature (T) (e.g., Mottana, 1986). Numerous synthetic experiments have been conducted at various fo, values controlled by buffers of certain mineral assem- blages (e.g., Abs-Wurmbach and Peters, 1999). Akasaka et al. (2003) conducted a series of hydrothermal experi- ments at 400-500 °C and 300 MPa to constrain the stabil- ity of Mn-rich pumpellyite. Those authors confirmed Mn- (b) bearing pumpellyite as the precursor to piemontite and that it is stable at lower temperatures and higher oxygen TheYatsushiroArea fugacity, as proposed by Akasaka et al. (1988). Reinecke OD70 (1986) investigated phase relations in Mn-Al-rich quartz- OT19 ites of the upper-blueschist facies from Evvia and Andros Y1811 Serpentinite Tobiishisub-unit islands, Greece, to schematically show the low-T stability Hak koishisub-unit Sedimentarycomplexes of sursassite-bearing assemblages from Evvia when com- Metabasaltic rock Metasiliceous rock OMetachert (this study) pared to spessartine-bearing assemblages from Andros. These previous studies were based on the premise Figure 1. (a) Distribution of the Kurosegawa belt in SW Japan and study area. (b) Geotectonic map of the Kurosegawa Belt in the that oxygen is a perfectly mobile component (Korzhin- Yatsushiro area, Kyushu, Japan (modified after Saito et al., 2010; ski, 1959) and therefore that cations (e.g., Fe and Mn) Sato et al., 2016). The okhotskite-bearing sample was collected can be freely oxidized or reduced in metamorphic rocks. from outcrop OT10 in the western part of the Hakoishi sub-unit, However, some recent studies have found that oxygen which occupies the western half of the Otao unit. Localities of behaves as an inert element in natural systems (Tumiati other Mn-bearing samples, i.e., OT19, KY1811, and OD70, are also shown. et al., 2015; Li et al., 2016). In such cases, the oxidation states of cations are instead fixed, and fo, is lowered as a dependent variable that describes bulk cation oxidation melange that is exposed over an area of 2 × 10 km2, oc- states. Thus the phase relations among Mn-rich minerals cupying the western half of the Otao unit (Saito et al., need to be reconsidered in terms of oxygen mobility. Fur- 2010; Sato et al., 2014; Fig. 1). The Hakoishi subunit con- thermore, during geochemical interaction in subduction sists mainly of serpentinites, lawsonite-bearing blue- zones, oxygen is bound to highly oxidized rocks and is schists (LBSs), metasiliceous rocks, and minor metagran- thus transported deep into the Earth by subduction. Re- ite and is separated from adjacent units by serpentinites fining the relevant phase relations should provide a clear (Saito et al., 2010). Saito et al. (2010) mapped areas of n , ssnd e metabasaltic rocks and metasiliceous rocks, but in the mantle (Thomson et al., 2016). s p o s e In this paper, okhotskite-bearing rocks from a new lated with each other and form a coherent block (Ibuki et locality from the Hakoishi sub-unit of the Kurosegawa al., 2008). Sato et al. (2016) reported that lawsonite + Na belt in Kyushu are described, including the chemical com- -s ru sd o si dnd + auxnd position and Raman spectra of the okhotskite. A provi- semblage in LBS of the eastern part of the study area sional petrogenetic grid is developed for Mn-bearing min- and that lawsonite + Na amphibole + Na pyroxene forms erals under a quartz-excess environment. We also discuss the predominant one in LBS of the western part. Such the behavior of oxygen and the significance of highly oxi- a systematic distribution of low-variance assemblages dized minerals, including okhotskite, in subduction zones. shows that the pumpellyite + Na pyroxene + chlorite + H2O = lawsonite + Na amphibole reaction is recorded in GEOLOGICAL BACKGROUND AND the study area, and an increase in grade to the west is MINERALOGY associated with hydration. The peak metamorphic condi- tions have been estimated to be 200-300 °C and 0.60-0.80 The Usuki-Yatsushiro Tectonic Line (UYTL) trends NE- GPa (Sato et al., 2016). Notably, the Hakoishi LBS (i.e., SW across Kyushu Island, Japan (Fig. 1). In the Yatsu- metabasalt) lacks epidote (Fujimoto et al., 2010; Kami- shiro area, the Kurosegawa Belt lies to the south of the mura et al., 2012; Sato et al., 2014; Sato et al., 2016). UYTL. The studied sample was collected from the west- Metasiliceous rocks in the study area are rich in ern part of the Hakoishi subunit, which is a serpentinite manganese and rare-earth elements (REEs), and are
A new occurrence of okhotskite and its petrological significance 433 a b layer Fe-Mn-rich layer Qtz veins Fe-Mn-rich cm 50um C (d) Okhotskite 30μm 100μm layer Figure 2. Texture of the okhotskite-bearing sample (OT10J). (a) Hand-specimen photograph. (b) Photomicrograph of okhotskite grains within (PPL). found in the abandoned Taneyama mine (Yoshimura, (Figs. 2a and 2d). The Fe-Mn-rich layers are fractured 1969; Yabuta and Hirajima, 2018). The Mn-rich rocks and include quartz-dominated veins that contain okhot- have been interpreted as impure metacherts that formed skite (Figs. 2a-2c). The different parts of the sample con- from siliceous oozes mixed with oceanic sediments, giv- sist of different mineral assemblages maybe reflecting the en that manganese and REEs are common in abyssal local “bulk-rock’ composition (Table 1). The quartz-rich sediments (Yoshimura, 1969; Togari et al., 1988; Kato layers consist of quartz + Na pyroxene + piemontite + et al., 2011). A variety of Mn-bearing minerals have been braunite + albite + apatite. The Fe-Mn-rich layers are reported from metasiliceous rocks related to the aban- composed of hematite + braunite + quartz ± albite ± apa- doned Taneyama mine (e.g., taneyamalite: Aoki et al., tite ± titanite. The quartz veins are composed predomi- 1981; howieite: Ibuki et al., 2008; Mn3+-bearing lawson- nantly of quartz along with Na amphibole + okhotskite ± ite: Ibuki et al., 2010). braunite ± albite ± apatite, and notably developed within The okhotskite-bearing sample (OT101/J) is a band- Fe-Mn-rich layers. ed rock composed of centimeter-thick intercalated quartz-rich and Fe-Mn-rich layers composed of fine- METHODS grained (<30 μm) minerals. Boundary layers (<100 μm thick) composed of Na pyroxene and amphibole (with The chemical composition of okhotskite was determined minor hematite, piemontite and okhotskite) occur at the using a Hitachi S3500H scanning electron microscope contacts between the quartz-rich and Fe-Mn-rich layers equipped with an energy-dispersive X-ray analyzer
434 W. Yabuta and T. Hirajima Table 1. Mineral assemblages observed in sample OT10 Qz Hem Okh Pmt Br Na-Amp Na-Px Ab Ttn Ap Qz-rich layers · · · + 十 Fe/Mn- Matrix · Rich layers Qz-veins · + Boundary layers + O, common; +, rare. (EDAX?) at Kyoto University, Kyoto, Japan. Operating Fe3+, V3+) = 8.00 (after Kato et al., 1981; Togari and Aka- conditions included an accelerating voltage of 20 kV, a saka, 1987). Although H2O content in the Hakoishi okhot- beam current of 500 pA, and a spot size of 5 μm. Fol- skite was not measured in the present study, it is calculated lowing materials were used as standards; albite for Na, as 7.15 wt% H2O by difference of the average chemical synthetic periclase for Mg, synthetic corundum for Al, composition, which is similar to the values (6.89 wt% synthetic quartz for Si, orthoclase for K, synthetic wol- H2O) of the Tokoro okhotskite determined by Togari and lastonite for Ca, synthetic rutile for Ti, vanadium metal Akasaka (1987). From the given formula for the average for V, manganese metal for Mn, and hematite for Fe. A composition of the Hakoishi okhotskite and the nomencla- ZAF correction scheme was used to process the X-ray ture scheme after Passaglia and Gottardi (1973), this min- intensity data. eral should be called okhotskite-(Mn2+). Raman spectra of okhotskite were acquired using Piemontite is present as a major Mn-bearing mineral an NRS-3100 JASCO spectrometer at the Department in the quartz-rich layers and boundary layers, and shows of Geology and Mineralogy, Kyoto University, using characteristic striking pinkish-red to yellow pleochroism a diode-pumped solid-state (DPSS) 532 nm laser. The (Fig. 2d). Some grains show oscillatory zoning, which is wave number was calibrated using the 520.7 cm-1 line attributed to variable REE contents (e.g., La and Nd). The produced by a Si-wafer, and the spectrum of Ne. chemical composition of the REE-poor zones is close to We also conducted thermodynamic calculations to that of ideal piemontite: Ca2Al2Mn3+SisO12(OH) (Table 2). estimate phase relations among Mn-bearing minerals. Na pyroxene occurs as acicular grains with a low Thermodynamic parameters of mineral phases followed molar fraction of diopside (Di < 15 mol%) (Table 2). those in the standard state of Holland and Powell (1998) The proportions of aegirine (Aeg) and jadeite (Jd) are or were calculated by the summation method of Fyfe et variable and reflect the local bulk-rock composition, with al. (1958) based on the database of Robie and Hemi- Aeg45JdsoDis in the quartz-rich layers and Aeg7oJd2oDi10 ngway (1995). Parameters of fuid species were calculat- in the boundary layers (Yabuta and Hirajima, 2018). The ed after Woolley (1953) and Burnham et al. (1969). Na amphibole is Mg-rich, consistent with a high ratio of Schreinemakers’ nets were generated with Calnet ver. ferric to ferrous iron, and is classified as magnesiorie- 2.1 developed by Yoshida and Hirajima (1999). beckite according to the scheme of Leake et al. (1997) Mineral abbreviations follow Whitney and Evans (Table 2). Na amphibole grains in boundary layers and (2010), except for okhotskite (Okh), braunite (Br), and quartz veins show nearly identical compositions. Hema- sursassite (Sus). tite and braunite occur in the Fe-Mn-rich layers, reflect- ing the highly oxidized Si-rich composition (Abs-Wurm- CHEMICAL ANALYSES AND bach, 1980). Braunite also occurs as tiny grains of ~ 20 RAMAN SPECTROSCOPY μm diameter in the quartz-rich layers. Figure 4 compares the Raman spectrum of an okhot- Okhotskite shows characteristic orange to dark-orange skite grain with those of related minerals. This is the first pleochroism, commonly occurs within the quartz veins as report of Raman spectrum of okhotskite, so the character- fine-grained (<30 μm long) crystals, and is uncommon in istic peaks are described. Peaks are observed consistently the boundary layers (Figs. 2b and 2c). The average chemi- close to 362, 480, 537, 563, 593, 684, and 921 cm-1. Even cal formula is (Ca7.62Mn2.16)≥7.78(Mn27Mg1.29)≥4.00(Mn4.13 though the Raman spectra of julgoldite-Fe and Mg-pum- Fe2.26A11.36V323Ti0.02)≥8.00Si11.86O4.02(OH)16.98, -with a Va- pellyite also show a peak near 550 cm-', the spectrum of riety in Fe:Mn ratio (Fetot/Mntot = 0.13-0.56; Table 2 and okhotskite show few similarities to those of other pum- Fig. 3). Total Fe was assumed as trivalent, and the oxida- pellyite group minerals. The peaks close to 362, 480, and tion state of the Mn was calculated based on Z(Mn3+, Al, 563 cm-1 are considered to be diagnostic of okhotskite.
A new occurrence of okhotskite and its petrological significance 435 Table 2. Mineral compositions for sample OT10 Okhotskite Pmt Na-Px Braunite Na-Amp Mineral (∑ cation = 32) (0 = 12.5) (0=6) (0 = 12) (0 = 23) Sample OT10J OT101 Mean OT10J 20 75 24 26 33 N=5 2 94 123 115 SiO2 33.77 33.85 32.69 32.63 33.35 33.26 36.63 55.84 11.44 57.60 TiO2 IPq IPq 0.07 IPq 0.31 0.10 IPq 0.32 0.10 0.48 Al2O3 3.58 2.17 3.93 3.82 2.72 3.24 19.85 11.69 0.33 1.39 V2O3 0.35 0.82 1.10 1.11 0.63 0.80 n.d. n.d. n.d. n.d. Fe2O3* 4.03 10.45 9.84 9.10 8.69 8.42 2.42 14.81 5.23 12.78 Mn2O3** 18.93 14.74 12.36 13.02 15.27 14.98 14.10 0.45 66.88 0.01 MnO** 10.59 8.91 8.70 9.48 11.24 9.68 0.53 IPq 10.14 1.20 MgO 1.31 3.02 3.41 2.90 1.52 2.43 0.20 1.80 0.24 15.52 CaO 19.75 19.59 20.53 20.05 19.79 19.94 22.17 1.51 1.73 1.48 CuO n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.95 n.d. Na2O n.d. n.d. n.d. n.d. n.d. n.d. n.d. 13.66 n.d. 7.29 K2O n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.15 92.31 93.55 92.63 92.11 93.52 92.85 95.90 100.08 97.04 97.90 Total Si 12.308 12.146 11.685 11.783 12.065 12.010 3.019 1.999 1.169 8.012 Ti 0.000 0.019 0.000 0.084 0.026 0.009 0.008 0.050 A1 1.538 0.918 1.656 1.626 1.160 1.381 1.928 0.493 0.040 0.228 V3+ 0.104 0.235 0.316 0.322 0.184 0.191 Fe3+ 1.106 2.821 2.646 2.473 2.366 2.289 0.150 0.399 0.403 1.338 Mn3+ 5.252 4.027 3.363 3.579 4.206 4.120 0.884 0.012 5.204 0.001 Mn2+ 3.268 2.708 2.634 2.899 3.445 2.960 0.037 0.878 0.142 Mg 0.712 1.615 1.817 1.561 0.820 1.309 0.025 0.096 0.037 3.218 Ca 7.712 7.531 7.863 7.757 7.671 7.716 1.957 0.058 0.189 0.221 Cu 0.073 Na 0.948 1.966 K 0.027 Total 32.000 32.000 32.000 32.000 32.000 32.000 8.000 4.014 8.000 15.201 * Total Fe as Fe2O3. ** Recalculated values. For okhotskite, Mn in the X- and Y-sites was assumed to be Mn2+ and Mn3, respectively, and the Y-site was flled with A13+, Mn3t, and Fe3+ in this order, following the method proposed by Akasaka et al. (1997). For piemontite, braunite, and Na-Amp, Mn2+ and Mn3+ were calculated based on the charge balance. Pmt, piemontite; Na-Px, Na pyroxene; Na-Amp, Na amphibole. Bixbyite-bearing Bixbyite-free PROVISIONAL SCHREINEMAKERS' NET (host-rocks) (host-rocks) Tokoro belt (1) Hakoishi this study) Sanbagawabelt(3) Reinecke (1986) performed thermodynamic calculations Okh veins Kamisugai mine Kamogawa mine O Sausar group(2) among Mn-bearing minerals, such as piemontite, Mn- pumpellyite, sursassite, spessartine, and chlorite, in a pseudo-quaternary system of Mg-Mn*(= Mn2+ + Mn3+)- Ca-Al, with excess Qz, H2O, and O2. This is applicable to C the Hakoishi metacherts, as quartz is saturated in the 0& △ whole sample and relevant minerals except in the pres- ence of garnet. 落 However, the Schreinemakers’ net proposed by Re- inecke (1986) has a fundamental shortcoming. The net n n ] =) d l, od ga s ： 紫 senting reactions where the two Ca-bearing phases are ** /Fe3+ Julgoldite Okhotskite Mn3 absent. As Reinecke's analyses do not deal with other Ca-phases, when one occurs as a reactant in a reaction, Figure 3. Y-site occupancy in okhotskite. Sources: (1) Tokoro the other is always present as a product. Consequently, Mn-Fe ore deposits (Togari et al., 1988). (2) Bichhua Forma- tion, Sausar Group, India (Dasgupta et al., 1991). (3) Mn ore either of the two is present on the same side as the invar- deposits, Sanbagawa belt (Minakawa, 1992). iant point [Pm, Pp] across a reaction curve, contravening
436 W. Yabuta and T. Hirajima Table 3. Thermodynamic parameters of Schreinemakers’ analyses Mineral Chemical composition Volume [J/bar·mol] Entropy [J/K·mol] Lawsonite (Lws) Ca(A11.8Mn3+0.2)Si2O7(OH)2:H2O 10.21 (1) 239.0 (2) Sursassite (Sus) (Cao.5Mn2+3.5)(A14Mn3+)Si6O21(OH)7 28.30 (3) 724.7(2) Piemontite (Pmt) Ca2Al2Mn3+Si3O12(OH) 14.02 (4) 307.7(2) Braunite (Br) (Mn2+0.9Ca0.1)Mn3+6SiO12 12.51(4) 415.8 (2) Spessartine (Sps) (Mn2+2.7Ca0.3)Al2Si3O12 11.89 (5) 314.1 (5) Ca4(Mn2+0.5Mg.5)(Mn33.5A1.5)Si6O21(OH)7 Okhotskite (Okh) 30.70 (6) 753.7(2) Chlorite (Chl) MgsAl2Si3O10(OH)8 21.09 (5) 410.5 (5) Quartz (Qz) SiO2 2.269 (5) 41.50 (5) Sources: () Ibuki et al. (2010). (2) calculated in this study (see the text). (3) Nagashima et al. (2009). (4) Anthony et al. (1995) and references therein. (5) Holland and Powell (1998). (6) Togari and Akasaka (1987). fied in the Hakoishi metacherts (Yabuta and Hirajima, （a) Okhotskite 2019), making it suitable for estimating phase relations in the metamorphic rocks of interest. In addition, compo- sitions of garnet, braunite, and sursassite were revised to contain Ca, reflecting natural occurrences in Greece and Kyushu (Table 3). The detail petrography and mineralogy of these Mn-bearing minerals in the Hakoishi sub-unit will be shown elsewhere. The Calnet program yields two possible bundles of reactions, corresponding to the invariant points of [Pmt] and [Okh]. We selected the latter to account for the min- h Julgoldite-(Fe) eral assemblages of Evvia-Andros (Reinecke, 1986). In the provisional petrogenetic grid (Fig. 5), the rep- resentative mineral assemblages of Evvia (Pmt + Sus + Br + Chl) and Andros (Pmt + Sps + Br + Chl) are stable around the invariant point [Lws]. The two assemblages (C) Pumpellyite-(Mg) are separated by the following reaction: (RRUFFDB) Sus + Br = Pmt + Sps + Chl + Qz + H2O + O2 (1) 500 1000 corresponding to the reaction (8) of Reinecke (1986). Raman shift (cm-1) In contrast, mineral assemblages in the Hakoishi metacherts (Sus + Chl + Pmt, Lws + Sus + Chl, and Figure 4. (a) Raman spectra of okhotskite. (b) and (c) Related Lws + Br + Chl; W. Yabuta, unpublished data, 2020) minerals. Julgoldite-Fe (ID R070725) and pumpellyite-Mg (R120172) are from the RRUFF database. are stable around the invariant point [Okh]. This does not contradict the occurrence of okhotskite, which can stably coexist with Br and Pmt on the low-P side of a basic rule of phase diagrams. the reaction curve: In addition, the petrogenetic grid contains viridine, Mn-bearing andalusite. Viridine has been reported from Lws + Okh + O2 = Pmt + Br + Chl + Qz + H2O many Mn-rich metamorphic rocks of relatively lower (2). metamorphic pressure but has not been identified on Ev- via or Andros islands (Greece) or in Hakoishi. Therefore, On the whole, the petrogenetic grid indicates that it is questionable to consider phase relations in blueschist the Evvia-Andros assemblages are stable at higher pres- facies (kyanite stability) with viridine. Therefore, in place sure and temperature compared with the Hakoishi meta- of viridine, we used Mn-bearing lawsonite (Ibuki et al., cherts. This is consistent with estimated P-T conditions 2010). The Mn-bearing lawsonite is ubiquitously identi- of the two areas: 350-450 °C and 1.0-1.2 GPa for south-
A new occurrence of okhotskite and its petrological significance 437 +Qtz Mg ite-(Mn2+) commonly coexist with okhotskite (Togari et +HO al., 1988; Dasgupta et al., 1991). This can be explained in +O2 terms of the higher metamorphic pressure that affected the Hakoishi subunit; i.e., 200-300 °C and 0.60-0.80 GPa as O [Br] AI4 estimated for Zone 2 at Hakoishi (Sato et al., 2016), and 200-230 °C and 0.25-0.35 GPa as estimated for rock units Sps [Chl] Mn* in the Kokuriki mine (Sakakibara, 1991). Instead of pum- pellyite-(Mn?+), sursassite and Mn-lawsonite occur as major minerals in the Hakoishi metacherts. Figure 3 compares the compositions of the okhot- Andros O[Lws] skite described in this study with those of previous stud- ies. A wide range in Mn:Al ratio but a narrow range in Fe:Mn ratio in okhotskite from the Tokoro Belt and Sau- Hakoishi Pmt+Sps+Br+CI sar Group have been documented by Togari et al. (1988) OD70 Evvia and Dasgupta et al. (1991), respectively. In contrast, the okhotskite described in the present study shows a narrow range in Mn:Al ratio and a wide range in Fe:Mn ratio (Fig. 3). Minakawa (1992) and Kassandrov and Mazurov Reinecke(1986) (2009) reported similarly uniform Mn:Al ratios and con- sistently low Fe2O3 contents in okhotskite from the San- P bagawa belt and Askiz ore district. OT19 KY1811 These patterns of Y-site occupancy may be related to the presence or absence of bixbyite in host rocks, whereby A13+-Mn3+ substitution is dominant in the for- T mer case and Fe3+-Mn3+ substitution is dominant in the latter case. As most bixbyite-free rocks contain quartz, Figure 5. A provisional Schreinemakers’ net indicating stability relations among okhotskite and associated Mn-bearing phases assemblages of braunite + quartz and braunite + bixbyite (following a revised methodology after Reinecke, 1986). The may translate into Si-saturation in highly oxidized sys- P-T ratio is arbitrary. tems (Abs-Wurmbach, 1980). However, further investi- gation is required, as some previous studies have reported ern Evvia (Shaked et al., 2000) and 200-300 °C and okhotskite as mono-mineralic veins traversing the Mn- 0.60-0.80 GPa for Hakoishi (Sato et al., 2016). With ref- pumpellyite veinlets with/without piemontite (Togari et erence to the study of Akasaka et al. (2003), who sug- al., 1988); for such cases, equilibria within host rocks gested that Mn-pumpellyite is stable below 500 °C at 300 are not assured. MPa, although schematic, the petrogenetic grid in Figure The provisional Schreinemakers’ net presented here 5 describes the phase relations below that temperature. (Fig. 5) predicts that okhotskite and Mn3+-bearing law- sonite are characteristically stable at higher pressure com- DISCUSSION pared with the P-T conditions determined for Hakoishi and Evvia-Andros. These minerals contain abundant In the studied sample, okhotskite occurrence is limited to Mn3+ and host substantial contents of H2O (Okh: 6.89 an environment closely associated with braunite and he- wt%, Togari and Akasaka, 1987 and 7.15 wt% in this matite, where Mn3+ is abundant. In contrast, quartz-rich study; Lws: 11.3 wt%, calculated from data of Ibuki et layers contain piemontite as the Ca-bearing mineral, and al., 2010). This clearly shows that siliceous sediments Al is relatively concentrated in the presence of albite. can contain high contents of water in the deep Earth These observations suggest that local chemical composi- when impurities such as Mn-rich oceanic sediments are tions strongly control the mineral assemblage, supporting present, although such sediments have generally been the finding of Akasaka et al. (1988) and Togari et al. thought to be water poor (Hacker, 2008). (1988) on the compositional variations in Mn-rich pum- In terms of oxygen transport, the Schreinemakers' pellyite and piemontite. net in Reinecke (1986) assumes perfect mobility of oxy- It is also important to note that no Mn-rich pum- gen (Korzhinski, 1959), whereby oxygen is released as a pellyite (with a high occupancy of Al in the Y-site) is fuid from oxidized solid phases. Under reduced mobility observed in the Hakoishi metacherts, although pumpelly- in a subduction zone, oxygen fuid is transported into the
438 W. Yabuta and T. Hirajima mantle wedge along with highly oxidized minerals (as ite, Mn2+Mn3+Og/SiO4, in the System Mn-Si-O at 1 atm in noted by Tumiati et al., 2015). Indeed, oxygen fuid is Air. Contributions to Mineralogy and Petrology, 71, 393-399. Abs-Wurmbach, I. and Peters, T. (1999) The Mn-Al-Si-O system: highly complicated when considering phase relations, ab experimental study of phase relations applied to paragen- as its high entropy sometimes leads to the prediction of eses in manganese-rich ores and rocks. European Journal of contradictory phase relations. Mineralogy, 11, 45-68. For instance, phase relations between pumpellyite Akasaka, M., Sakakibara, M. and Togari, K. (1988) Piemontite from the manganiferous hematite ore deposits in the Tokoro and epidote groups are discussed as simplified oxidation Belt, Hokkaido, Japan. Mineralogy and Petrology, 38, 105- reactions (e.g., Fe-Pmp + O2 = Ep + H2O; Schiffman and 116. Liou, 1983; Mn-Pmp + O2 = Pmt + H2O; Akasaka et al., Akasaka, M., Kimura, Y., Omori, Y., Sakakibara, M., et al. (1997) 1988, 2003). These reactions, however, would have neg- 57Fe Mossbauer study of pumpellyite-okhotskite-julgoldite ative slope (△S < 0 and △V > 0, owing to high entropy of series minerals. Mineralogy and Petrology, 61, 181-198. Akasaka, M., Suzuki, Y. and Watanabe, H. (2003) Hydrothermal O2-fluid) with pumpellyite group stable at higher P-T Synthesis of Pumpellyite-Okhotskite Series Minerals. Miner- side, even when some other solid phases are present. alogy Petrology, 77, 25-37. Needless to say, such reaction curves on P-T field would Anthony, J.W., Bideaux, R.A., Bladh, K.W. and Nichols, M.C. not explain the high-temperature stability of the epidote (1995) Handbook of Mineralogy volume Il. Tucson, AZ: Mineral Data Publishing, ISBN 9780962209703. group against pumpellyite group in metabasites (e.g., Aoki, Y., Akasako, H. and Ishida, K. (1981) Taneyamalite, a new Nakajima et al., 1977). manganese silicate mineral from the Taneyama mine, Kuma- Such a simple case is sufficient to demonstrate prob- moto Prefecture, Japan. Mineralogical Journal, 10, 385-395. lems in oxygen-mobile models. The effect of oxygen Brigatti, M.F., Caprilli, E. and Marchesini, M. (2006) Poppite, the V3+ end-member of the pumpellyite group: Description and fluid on phase relations will be discussed in detail else- crystal structure. American Mineralogist, 91, 584-588. where. In the present study, it is noted that highly oxi- Burnham, C.W., Holloway, J.R. and Davis, N.F. (1969) Thermo- dized minerals (e.g., Okh and Br) are stable at high P-T dynamic properties of water to 1,000 °℃ and 10,000 bars. even if oxygen is released freely from the system as a Memoirs Geological Society of America, Special Paper 132. perfectly mobile element. This indicates that okhotskite Dasgupta, S., Chakraborti, S., Sengupta, P., Bhattacharya, P.K., serves as a component of the oxygen budget in deep sub- et al. (1991) Manganese-rich minerals of the pumpellyite duction zones, in addition to its role as a water reservoir group from the Precambrian Sausar Group, India. American Mineralogist, 76, 241-245. (Togari and Akasaka, 1987; Togari et al., 1988). Howev- Fujimoto, Y., Kono, Y., Hirajima, T., Ishikawa, M. and Arima, M. er, further investigation is required regarding the fate of (2010) P-wave velocity and anisotropy of lawsonite and epi- highly oxidized Mn-bearing minerals, including okhot- dote blueschists: Constraints on wave transportation along skite, in subduction environments, where their oxidizing subducting oceanic crust. Physics of the Earth and Planetary Interiors, 183, 219-228. capacity may oxidize the mantle hanging wall. Fyfe, W.S., Turner, F.J. and Verhoogen, J. (1958) Metamorphic reactions and metamorphic facies. Memoirs Geological Soci- ACKNOWLEDGMENTS ety of America, 73. 21 51. Hacker, B.R. (2008) H2O subduction beyond arcs. Geochemistry, We sincerely thank Dr. M. Akasaka, Professor Emeritus Geophysics, Geosystems, 9， Q03001, DOI:10.1029/ 2007GC001707. of Shimane University, and Dr. S. Endo for their sincere Holland, T.J.B. and Powell, R. (1998) An internally consistent ther- and constructive comments, which guided the revision of modynamic data set for phases of petrological interest. Jour- an earlier draft of the manuscript. We also thank Profes- nal of Metamorphic Geology, 16, 309-343. sor M. Satish Kumar for his efficient editorial handling of Ibuki, M., Fujimoto, Y., Takaya, M., Miyake, A. and Hirajima, T. (2008) Howieite in meta-manganese siliceous rocks of Kuro- the manuscript. We are grateful for the technical advice segawa belt, western Kyushu, Japan. Journal of Mineralogical offered and the high-quality thin-section samples pre- and Petrological Sciences, 103, 365-370. pared by Mr. M. Takaya and the revision of the Calnet Ibuki, M., Ohi, S., Tsuchiyama, A. and Hirajima, T. (2010) Analy- program by Mr. R. Kato. We acknowledge students of sis of Mn-bearing lawsonite occurring in meta-siliceous rocks Matsubase High School, Kumamoto, for their contribu- in Hakoishi serpentinite melange of Kurosegawa Belt, Central tion to the first report of piemontite in the study area. Kyushu, Japan. Journal of Mineralogical and Petrological Sci- ences, 105, 340-345. This work was supported by JSPS KAKENHI Grant No. Ivanov, O.K., Arkhangel'skaya, V.A., Miroshnikova, L.O. and Shi- JP19H1999 to T.H. lova, T.A. (1981) Shuiskite, the chromium analog of pum- pellyite, from the Bisersk deposit, Urals. Zapiski Vserossiy- skogo Mineralogicheskogo Obshchestva, 110, 508-512 (in REFERENCES Russian). Kamimura, K., Hirajima, T. and Fujimoto, Y. 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(2011) of Mineralogical and Petrological Sciences, 109, 258-270. Deep sea mud in the Pacific Ocean as a potential resource Sato, E., Hirajima, T., Yoshida, K., Kamimura, K. and Fujimoto, Y. for rare earth elements. Nature Geoscience, 4, 535-539. (2016) Phase relations of lawsonite-blueschists and their role Korzhinski, D.S. (1959) Physicochemical Basis of the Analysis of as a water-budget monitor: a case study from the Hakoishi the Paragenesis of Minerals. Consultants bureau, New York. sub-unit of the Kurosegawa belt, SW Japan. European Journal Leake, B.E., Woolley, A.R., Arps, C.E.S., Birch, W.D., et al. of Mineralogy, 28, 1029-1046. (1997) Nomenclature of Amphiboles: Report of the Subcom- Schiffman, P. and Liou, J.G. (1983) Synthesis of Fe-pumpellyite mittee on Amphiboles of the International Mineralogical As- and its stability relations with epidote. Journal of Metamor- sociation, Commission on New Minerals and Mineral Names. phic Geology, 1, 91-101. Canadian Mineralogist, 35, 219-246. Shaked, Y., Avigad, D. and Garfunkel, Z. (2000) Alpine high-pres- Li, J., Gao, J., Klemd, R., John, T. and Wang, X. (2016) Redox sure metamorphism at the Almyropotamos window (southern processes in subducting oceanic crust recorded by sulphide- Evvia, Greece). Geological Magazine, 137, 367-380. bearing high-pressure rocks and veins (SW Tianshan, China). Thomson, A.R., Walter, M.J., Kohn, S.C. and Brooker, R.A. Contributions to Mineralogy and Petrology, 171, 72. (2016) Slab melting as a barrier to deep carbon subduction. Minakawa, T. (1992) Study on characteristic mineral assemblages Nature, 529, 76-83. and formation process of metamorphosed manganese ore de- Togari, K. and Akasaka, M. (1987) Okhotskite, a new mineral, an posits in the Sanbagawa belt. Memoirs of the Faculty of Sci- Mn3-dominant member of the pumpellyite group, from the ence, Ehime University, 1, 1-74. (in Japanese with English Kokuriki mine, Hokkaido, Japan. Mineralogy Magazine, 51, abstract) 611-614. Minakawa, T., Fukushima, H., Nishio-Hamane, D. and Miura, H. Togari, K., Akasaka, M., Sakakibara, M. and Watanabe, T. (1988) (2008) Epidote-(Sr), CaSrAlFe3+(SizO7)(SiO4)(OH), a new Mineralogy of manganiferous iron ore deposits and chert from mineral from the Ananai mine, Kochi Prefecture, Japan. Jour- the Tokoro Belt, Hokkaido. Mining Geology Special Issue, nal of Mineralogical and Petrological Sciences, 103, 400-406. 12, 115-126. Moore, P.B. (1971) Julgoldite, the Fe2+-Fe3+ dominant pumpelly- Tumiati, S., Godard, G., Martin, S., Malaspina, N. and Poli, S. ite. Lithos, 4, 93-99. (2015) Ultra-oxidized rocks in subduction melanges? Decou- Mottana, A. (1986) Blueschist-facies metamorphism of manganif- pling between oxygen fugacity and oxygen availability in a erous cherts: A review of the alpine occurrences. Geological Mn-rich metasomatic environment. Lithos, 226, 116-130. Society of America Memoir, 164, 267-299. Whitney, D.L. and Evans, B.W. (2010) Abbreviations for names of Nagashima, M, Akasaka, M., Minakawa, T., Libowitzky, E. and rock-forming minerals. American Mineralogist, 95, 185-187. Armbruster, T. (2009) Sursassite: Hydrogen bonding, cation Woolley, H.W. (1953) Thermodynamic Properties of Molecular order, and pumpellyite intergrowth. American Mineralogist, Oxygen. National Bureau of Standards, Washington, D.C. 94, 1440-1449. Yabuta, W. and Hirajima, T. (2018) Mineralogy of Highly-Oxi- Nakajima, T., Banno, S. and Suzuki, T. (1977) Reactions leading to dized Lawsonite-Blueschist Facies Metachert from the Hako- the disappearance of pumpellyite in low-grade metamorphic ishi sub-unit, the Kurosegawa Belt, Kyushu and Proposal for rocks of the Sanbagawa metamorphic belt in central Shikoku, the Piedmontite Forming Reaction. Abstract of 2018 JpGU Japan. Journal of Petrology, 18, 263-284. Annual meeting, SMP-PO2, Chiba, Japan. Palache, C. and Vassar, H.E. (1925) Some minerals of the Kewee- Yabuta, W. and Hirajima, T. (2019) Blueschist facies impure meta- nawan copper deposits: Pumpellyite a new mineral; sericite; chert as remarkable oxygen- and overlooked water-carriers in saponite. American Mineralogist, 10, 412-418. the subduction zone. Abstract of 2019 Annual Meeting of Ja- Passaglia, E. and Gottardi, G. (1973) Crystal chemistry and no- pan Association of Mineralogical Sciences, R8-08, Fukuoka, menclature of pumpellyites and julgoldites. Canadian Miner- Japan. alogist, 12, 219-223. Yoshida, D. and Hirajima, T. (1999) The algorithm for drawing Reinecke, T. (1986) Phase relationships of sursassite and other petrogenetic grids with Schreinemakers’ method. Journal of Mn-silicates in highly oxidized low-grade, high-pressure Mineralogy, Petrology and Economic Geology, 94, 254-260. metamorphic rocks from Evvia and Andros Islands, Greece. Yoshimura, T. (1969) Supplement to “Manganese Ore Deposits of Contributions to Mineralogy and Petrology, 94, 110-126. Japan", Part I. Science Reports, Department of Geology, Robie, R.A. and Hemingway, B.S. (1995) Thermodynamic Proper- Kyushu University, 9 (Special Issue No. 2), 902 (in Japanese). ties of Minerals and Related Substances at 298.15 K and 1 Bar (10 Pascals) Pressure and at Higher Temparatures. U.S. Geo- logical Survey Bulletin, 2131. Manuscript received August 31, 2019 Saito, M., Takarada, S., Toshimitsu, S., Mizuno, K. et al. (2010) Manuscript accepted August 5, 2020 Geological map of Japan 1:200,000 Yatsushiro and a part of PublishedonlineOctober28,2020 Nomo Zaki. Geological Survey of Japan, AIST (in Japanese Manuscript handled by M. Satish-Kumar
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Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 CrossMark DOI 10.1007/s00531-016-1437-6 ORIGINAL PAPER Early Paleozoic subduction initiation volcanism of the Iwatsubodani Formation, Hida Gaien belt, Southwest Japan Kazuhiro Tsukada1 · Koshi Yamamoto2 · Onon Gantumur² · Manchuk Nuramkhaan? Received: 31 March 2016 / Accepted: 14 December 2016 / Published online: 24 January 2017 ? Springer-Verlag Berlin Heidelberg 2017 Abstract In placing Japanese tectonics in an Asian con both these ophiolites and the Iwatsubodani Formation text, variation in the Paleozoic geological environment is probably coexisted in a primitive SSZ system in the early a significant issue. This paper investigates the geochemis- Paleozoic. try of the lower Paleozoic basalt formation (Iwatsubodani Formation) in the Hida Gaien belt, Japan, to consider its Keywords Volcanic arc at SSZ · Fore-arc basalt (FAB) · tectonic setting. This formation includes the following two Hitoegane succession · Fukuji area (Gifu Pref.) types of rock in ascending order: basalt A with sub-ophitic texture and basalt B with porphyritic texture. Basalt A has a high and uniform FeO*/MgO ratio, moderate TiO2, high Introduction V, and low Ti/V. The HFSE and REE are nearly the same as those in MORB, and all the data points to basalt A being The geology of the East Asian continental margin devel. the “MORB-like fore-arc tholeiitic basalt (FAB)” reported, oped through various processes, such as accretion, col- for example, from the Izu-Bonin-Mariana arc. By contrast, lision, and strike-slip tectonic movement. Accretion of basalt B has a low FeO*/MgO ratio, low TiO2, and low V trench-fill sediments and oceanic plate cover, as well as and Ti/V. It has an LREE-enriched trend and a distinct neg- collision of terranes with strike-slip tectonic movement, ative Nb anomaly in the MORB-normalized multi-element created the basic geotectonic framework of southwest- pattern and a moderately high LREE/HREE. All these fac- ern (Sw) Japan. Features of the accretionary complexes tors suggest that basalt B is calc-alkaline basalt. It is known of this area have been studied, and various tectonic mod- that FAB is erupted at the earliest stage of arc formation- els for their formation presented (e.g., Isozaki et al. 1990; Wakita 1989; Wakita et al. 1992). Although some sugges- namely, subduction initiation—and that boninitic/tholeiitic/ calc-alkaline volcanism follows at the supra-subduction tions have been proposed about the collision process and zone (SSZ). Thus, the occurrence of basalts A (FAB) and the role of strike-slip tectonic movement in the evolution B (calc-alkaline rock) is strong evidence of early Paleozoic of Sw Japan, a definitive answer remains a matter to be arc-formation initiation at an SSZ. Evidence for an early further researched (e.g., Charvet 2012; Faure and Char- Paleozoic SSZ arc is also recognized from the Oeyama, vet 1987; Faure et al. 1986; Tashiro 1994; Tazawa 2004; Hayachine-Miyamori, and Sergeevka ophiolites. Hence, Yamakita and Otoh 2000). The Hida Gaien belt, a major tectonic belt of SW Japan, contains fault-bound blocks of Paleozoic shelf facies rocks. It was formed through Meso- Kazuhiro Tsukada zoic strike-slip movements between the Hida and the San- tsukada@ num.nagoya-u.ac.jp gun-Renge belt (Otoh et al. 2003; Tsukada 2003, Fig. 1). To reveal the Paleozoic geodynamics, understanding the The Nagoya University Museum, Nagoya 464-8601, Japan Graduate School of Environmental Studies, Nagoya tectonic origin and environmental variation of each block 2 University, Nagoya 464-8601, Japan is important. The Hida Gaien belt exposes well-preserved School of Geology and Mining, Mongolian University Paleozoic successions and is one of the most important of Science and Technology, Ulaanbaatar, Mongolia sites for Paleozoic stratigraphy in Japan. Many geologists Springer 1430 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 b 50 km Russia Sea of Japan Sergeevka Hida belt China Hida Gaien belt N. Korea b Fig.3 Oeyama HidaMt S.Korea ophiolite Sangun-Renge belt Japan (with Oeyama ophiolite) Mino- Akiyoshi and Tamba belt Maizuru belts Hida Gaien belt Ultra-Tambabelt Fig. 1 Index map of the study area and paleontologists have studied the post-Silurian rocks Fukuji succession in this belt (e.g., Adachi 1985; Horikoshi et al. 1987; Igo 1956; Igo and Adachi 1981; Kamei 1952; Kato 1959; Nii- kawa 1980; Okazaki 1974; Tsukada and Takahashi 2000), but the geological environment in pre-Devonian times has not previously been investigated. The geochemistry of mafic volcanic rocks gives evidence for the tectonic setting of the volcanic activity, because the chemical composition niferous of the present mafic volcanic rocks varies according to their tectonic origins and this provides a basis for geodynamic Carb interpretation. Although little attention has so far been given to the geochemistry of the mafic volcanic rocks of the Iwatsubodani Formation of the Hitoegane succession, onian which includes the oldest fossil-bearing sedimentary strata in Japan, in the Hida Gaien belt (Fig. 2), it is critical to our understanding of the lower Paleozoic setting in Japan. This paper describes the lithology, stratigraphy, and geochem- istry of the Iwatsubodani Formation. Environmental vari- ation of the Hitoegane succession and its regional correla- tion is also discussed. Geological framework of the Hida Gaien belt Legend Southwest Japan is divided into the Inner and Outer zones, Mafic pyroclastic rocks and the former is composed of several tectonic belts. These are, from north to south: the Hida, Hida Gaien, Nagato, Sangun-Renge (including Oeyama ophiolite and Sangun-Renge metamorphic rocks), Akiyoshi, Maizuru, Fig. 2 Schematic stratigraphies of the Hitoegane, Fukuji, and Mor- Ultra-Tamba, Mino-Tamba, and Ryoke belts (Wakita 1989; ibu successions in the Hida Gaien belt (modified from Tsukada 2003) Wakita et al. 1992, Fig. 1). The Hida belt consists mainly of Paleozoic metamorphic rocks with Mesozoic cover, while mainly of lower Paleozoic ophiolite (Oeyama ophiolite), the Ryoke belt has a pre-metamorphic affinity with the upper Paleozoic blueschist-facies metamorphic rocks (San- Mino-Tamba belt. The Sangun-Renge belt is composed gun-Renge metamorphic rocks), and Mesozoic clastic Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1431 rocks. The Akiyoshi and Ultra-Tamba belts are Permian to tuffaceous clastic rocks and clastic rocks, Devonian lime- Triassic accretionary complexes, and the Mino-Tamba belt stone, Carboniferous limestone, and Permian felsic tuf- is upper Triassic to lower Cretaceous. The Maizuru belt faceous clastic rocks, clastic rocks, and mafic pyroclastic is an upper Paleozoic island arc entity covered by upper rocks; and (3) the Moribu succession composed of Devo- Paleozoic to upper Mesozoic strata. The Nagato belt is nian felsic tuff and tuffaceous clastic rocks, Carboniferous considered to be the western extension of the Hida Gaien mafic pyroclastic rocks, and Permian clastic rocks (Ehiro belt. The Sangun-Renge, Akiyoshi, Maizuru, and Ultra- et al. 2016; Tsukada 2003; Tsukada et al. 2004, Fig. 2). Tamba belts are widely exposed in the western part of SW The mafic volcanic rocks, examined in this paper, are Japan (Chugoku Mountains), while these belts are mostly assumed to be in fault contact with the Ordovician rocks lacking in the eastern part of SW Japan (Hida Mountains), in the Hitoegane succession, but the former was originally where the Hida Gaien belt is narrowly distributed between overlain by the latter as will be discussed later (Tsukada the Hida and Mino-Tamba belts (e.g., Tsukada et al. 2004, 1997, Fig. 3). The felsic tuff and tuffaceous clastic rocks of Fig. 1b). the Hitoegane succession, including Ludlowian to Pridol- The Hida Gaien belt, which is a significant tectonic zone ian radiolarians, trilobites, and corals, are correlated with in Japan, was formed as a result of Jurassic dextral and Cre- those of the Fukuji succession (Ehiro et al. 2016; Manchuk taceous sinistral strike-slip movements (Tsukada 2003). It et al. 2013b and references therein, Fig. 2). The Silurian includes several tectonic blocks of Paleozoic rocks which rocks of the Fukuji succession are largely in fault contact can be divided into the following three types based on with the surrounding formations, but are partly unconform- their lithostratigraphy: (1) the Hitoegane succession com- ably overlain by Devonian limestone (Igo 1990). Abun- posed of mafic volcanic rocks, and Ordovician to Devonian dant fossils, such as corals, stromatoporids, brachiopods, felsic tuff, tuffaceous clastic rocks, and clastic rocks; (2) trilobites, bivalves, ostracods, conodonts, and others, were the Fukuji succession composed mainly of Silurian felsic recorded from the limestone of early to middle Devonian Cenozoic Mafic volcanic Mafic intrusive rock volcanic rocks Permian mafic pyroclastic Ordovician- Silurian felsic tuff, Granite Moribu rocks andlava felsic tuffaceous clastic rocks Permian felsic tuffaceous and clastic rocks succession maficpyroclastic clastic rocks and clastic rocks Mafic pyroclastic rocks and lava Carboniferous limestone Gamata Crystalline Schists Silurianfelsic tuff,felsic Serpentinite clastic rocks Rocks of the Mino-Tamba belt Hitoegane Gaien Strike and dip of succession ddingplane Hida Fig. n! succession Mino Tambabelt 1 km Tsukada and Takahashi 2000) Springer 1432 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 age (Tsukada 2005 and references therein). Investigation pyroclastic rock has northwestward-up graded bedding of the biostratigraphy of foraminifers and corals from the striking N 65° E and dipping 90° near the boundary with Carboniferous formation of the Fukuji succession indicates the Hitoegane Formation (Figs. 3, 4, 5a, b). The pyroclas- an age of Visean to Gzhelian (e.g., Adachi 1985; Niikawa tic rock overlies the lava with an obscure boundary in the 1980; Watanabe 1991). This formation intercalates red and upper reaches of the valley. The bedding plane of the pyro- gray layers of shale and marl which are considered to have clastic rock near the boundary with the lava strikes N 309 been formed in fresh water lake conditions by deposition E and dips 90° (Fig. 4). The lava is generally fine-grained of paleosol produced by chemical weathering of limestone with a sub-ophitic texture, and the majority of the plagio- (Igo 1960). Smaller foraminifers and corals suggest a tropi- clase and clinopyroxene crystals are idiomorphic, fresh, cal to subtropical environment (Adachi 1985; Kato 1959). and clear (Fig. 5c). The lava partly shows an ophitic texture Devonian and Carboniferous formations are in fault con- with idiomorphic to hypidiomorphic laths of plagioclase tact each other (Fig. 3). They are, however, likely to have more than 1 mm in major axis. The lava was partly altered formed a superposing succession (Tsukada and Takahashi to the extent that some minerals have been replaced by sec- 2000; Fig. 2). The lower to upper middle Permian felsic ondary minerals, such as chlorite, muscovite, and opaque tuffaceous clastic rocks and clastic rocks in this succession minerals. The pyroclastic rock includes abundant angular are conformably overlain by mafic pyroclastic rock yielding clasts of basalt, from several millimeters to several tens of a middle Permian fusulinoidean fauna (Tsukada et al. 1999; centimeters in diameter (Fig. 5d). The clasts are divided Tsukada and Takahashi 2000; Figs. 2, 3). into two types by their texture. One type has a porphyritic The Devonian formation in the Moribu succession, texture with phenocrysts several millimeters in major axis which yield crinoids, tabulate corals, and plant fossils (Fig. 5e); the other shows a sub-ophitic texture similar to (Tazawa et al. 1997, 2000b), are in fault contact with other the lava. The majority of the clast is porphyritic basalt, and Paleozoic formations. The upper Visean to upper Kasimo- sub-ophitic basalt is rare. Larger phenocrysts of idiomor- vian or Gzhelian (?) mafic pyroclastic rock, which con- phic plagioclase and clinopyroxene are scattered in smaller tains corals, ammonoids, foraminifers, brachiopods, and interstitial patches of plagioclase in the porphyritic basalt trilobites (Tazawa et al. 2000a and references therein), is (Fig. 5e). The clasts with sub-ophitic texture include abun- unconformably overlain by middle Permian clastic rocks dant small needles of plagioclase less than 0.5 mm in major (Horikoshi et al. 1987; Niwa et al. 2004; Fig. 2). The clastic axis, in a groundmass. The minerals of the groundmass rocks yield brachiopods that form a mixed fauna of Boreal have been partly altered to chlorite, opaque minerals, and and Tethyan species having a close affinity to the fauna of others. In this paper, basalts with a sub-ophitic texture are inner Mongolia (e.g., Horikoshi et al. 1987; Tazawa 1991). referred to as basalt A and basalt clasts of the pyroclastic A boreal fusulinoidean, genus Monodiexodina, was found rock with porphyritic texture as basalt B. from this clastic rock formation (Niwa et al. 2004; Tazawa The Hitoegane Formation consists largely of alternat- et al. 1993). ing beds of felsic tuff, tuffaceous clastic rocks, and clas- tic rocks, well-bedded in beds from several millimeters to Geological description of the Hitoegane succession 1 m thick. The lower part of this formation includes clasts of mafic volcanic rock, and intercalates mafic pyroclastic The rocks of the Hitoegane succession in and around the rock layers, up to 20 m thick (Tsukada 1997). The clasts of Hitoegane area are divided into the following three forma- mafic volcanic rock in the Hitoegane Formation are several tions: (1) mafic volcanic rocks of the Iwatsubodani Forma- centimeters in diameter. The pyroclastic rock layers include tion; (2) the Ordovician to Devonian Hitoegane Formation abundant clinopyroxene crystals and two types of angular consisting of alternating beds of felsic tuff, tuffaceous clas- mafic volcanic rock clasts which appear in the mafic pyro- tic rocks, and clastic rocks with minor amounts of limestone clastic rock of the Iwatsubodani Formation (Tsukada 1997). lenses; and (3) Gamata Crystalline Schists (Tsukada 1997; The felsic tuff of the lower part of this formation yields Fig. 3). The inference is that the Iwatsubodani Formation is Ordovician conodonts, and zircon ages of ca. 436 Ma and in fault contact with the Hitoegane Formation, but the for- ca. 472 Ma (Nakama et al. 2010; Tsukada and Koike 1997; mer was originally overlain by the latter, as discussed later Fig. 4). The beds striking N 55°-60° E and dipping 60° (Tsukada 1997; Fig. 3). The Gamata Crystalline Schists are N-90° have northward-up sedimentary structures at the in fault contact with the Hitoegane Formation. The Iwat- lower part of the Hitoegane Formation. Minor amounts of subodani Formation, composed of mafic lava and pyroclas- fossiliferous limestone lenses are included in the upper part tic rock, is narrowly exposed along the Iwatsubodani val- of this formation. Silurian trilobites, corals, and radiolar- ley of this area (Figs. 3, 4). The lava exposed in the upper ians were reported from the felsic tuff and limestone of reach of the valley, striking N 30° E and steeply dipping the upper part of this formation (Kurihara 2007; Manchuk northward, shows eastward-up pillow structure, and the et al. 2013b and references therein). The mafic intrusive Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1433 Z 1300m 3070 1300m 09101-10_ Hitoegane Formation Iwatsubodani Valley Legend Iwatsubodani Cenozoicvolcanicrocks Mafic pyroclastic rocks Mafic lava Formation Mafic intrusive rock Felsic tuff, felsic tuffaceous clastic altA Fault 500 m altB Inferred Fault Strike： and dip of bedding plane Dam Fig.4 Route map of the study area showing the sampling locali- open circle ca. 472 Ma zircon locality, black star ca. 280 Ma zircon ties (modified from Tsukada 1997). C: Ordovician conondont local- locality (Nakama et al. 2010). White diamond is the locality of the ity (Tsukada and Koike 1997), white star ca. 436 Ma zircon locality, high-Mg andesite and boninite rock intrudes into the Iwatsubodani and Hitoegane Forma- (Nakama et al. 2010; Fig. 4). The intermediate dikes, gen- tions with an altered porous chilled margin (Tsukada 1997). erally less than 3 m wide, intrude into the Iwatsubodani and The rocks in the upper part of Hitoegane Formation gen- Hitoegane Formations and the mafic intrusive rock with a erally yield well-preserved radiolarian fossils, but those in clear chilled margin. The dike is composed mainly of idi- the lower part are metamorphosed around the mafic intru- omorphic plagioclase phenocrysts and a groundmass of sive rock and yield only recrystallized radiolarians which small needles of plagioclase. All these are unconformably are not identifiable. The intrusive rock is composed mainly covered by Cenozoic volcanic rocks (Fig. 3). of gabbro/diorite, and besides includes dolerite, basalt, and andesite as fine-grained facies. The gabbro/diorite is equi- granular, and is composed mainly of idiomorphic clinopy- Whole-rock chemistry of the basalts roxene, hornblende, and plagioclase (Figs. 3, 4). In a fine- of the Iwatsubodani Formation grained facies, idiomorphic phenocrysts of clinopyroxene, hornblende, and plagioclase, more than 1 mm in length, Sixteen samples from the Iwatsubodani Formation—six of lie in a groundmass of smaller plagioclase laths, granular basalt A and ten of basalt B—were analyzed for major and clinopyroxene, and interstitial chloritic material (Fig. 5f). trace elements, including rare-earth elements (REE). Zircon in the gabbro/diorite gives a U-Pb age of ca. 280 Ma Springer 1434 Int JEarth Sci(Geol Rundsch)(2017)106:1429-1451 Fig. 5 Photographs of the rocks in the study area. a Mafic lava showing pillow structure. It is up entirely composed of basalt A. b Polished surface of the mafic pyroclastic rock. Clear graded bedding is observed. c Photo- micograph of the basalt A from the pillow lava showing typical sub-ophitic texture. Plane polar- ized light. d Photomicograph of the mafic pyroclastic rock. It contains only volcanic material, such as volcanic rock fragments and crystals of clinopyroxene and plagioclase. rf rock frag- ment, Cpx clinopyroxene. Plane Lm polarized light. e Photomico- graph of the basalt B from the clast of the pyroclastic rock. Large clinopyroxene (Cpx) and plagioclase (Pl) phenocrysts are embedded in a finer matrix. Plane polarized light. f Photomicograph of the mafic intrusive rock with phenocrysts of clinopyroxene (Cpx), plagio- clase (Pl), and hornblende (Hb). Plane polarized light mm mm 1mm After coarse crushing, veins and altered parts of the major elements was estimated to be <1% for Si and about samples were excluded using x20 hand loupe. Each analy- 3% for other elements, except for TiO2 and MnO, whose sis sample weighed ca. 200 g. Major element compositions analytical precision is >3% when the measured level is as well as Co, Zn, Ga, and Zr were determined by X-ray <0.1% (Takebe and Yamamoto 2003) and that for trace ele- fluorescence (XRF; Rigaku Primus II ZSX equipped with ments (Co, Zn, Ga, and Zr) was estimated to be less than Rh X-ray tube, 50 kV, 60 mA), and other trace elements 10% (Yamamoto and Morishita 1997). The other trace ele- (V, Cr, Ni, Cu, Rb, Sr, Y, Nb, Mo, Pb, Ba, Hf, and Ta) and ments and REE were determined by a method based on that REE were analyzed by Inductively Coupled Plasma Mass described by Yamamoto et al. (2005) using ICP-MS. About Spectrometry (Quadrupole type ICP-MS; Agilent 7700x 30 mg of each sample was digested with a mixed solution with collision cell of He) installed at Nagoya University. of HF-HClO4 (2:1 by volume) at 150°C. After complete In the XRF analysis, glass beads were prepared by fus- evaporation of the acids, 2 ml of 1.7 N-HCl was added to ing mixtures of 1.5 g of powdered sample with 6.0 g of dissolve the cake. The residue was separated by centrifuga- lithium tetraborate. Calibration was carried out using the tion at 12,000 rpm with a 2 ml polypropylene tube. After standard rock samples issued by the Geological Survey centrifugation, the supernatant was transferred to another of Japan (GSJ) and the composite standards prepared by 10 ml Teflon beaker. The residue was then fused with HF- Yamamoto and Morishita (1997). Analytical precision of HClO4 (2:1 by volume) again at 150°C. The fused cake Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1435 was dissolved with about 2 ml 1.7 N-HCl by mild heating, wt%), Al2O3 (ca. 13 wt%), Fe2O3 (ca. 11 wt%), MnO (ca. and the solution was centrifuged at 12,000 rpm. In most 0.18 wt%), MgO (ca. 8.5 wt%), V (ca. 280 ppm), Co (ca. ) 1 pe (dd 5h e) e ‘(dd 9g ) uz dd 2 unnus soe pzuoo sm anpisa ou s aloclastitemassive flow, and metamorphosed mary of the latter paper is given below, dolerite dikes are intruded the gabbroic and sensitivities; the indium and Bi concentrations were mostly trend, correlated with the increase in SiO2 (Fig. 6). Basalt B the same throughout the analysis. The oxide generation fac- has lower FeO*/MgO ratio ranging from 0.8 to 1.9 and has represent feeder-dikes for the extrusive bas- the gabbroic rocks. The total thickness of basaltic rocks, but typical sheeted dike com- ples which were prepared from GSJ JB-1a and USGS BCR- Ta, P, Zr, Hf, Ti, Y, and HREE (Fig. 7a). The chondrite- 1. In the ICP-MS analysis, the correlation coefficients (R normalized REE patterns show a reducing trend—enriched values) of each element, calculated for five standard sam- in LREE and depleted in HREE (Fig. 7b). ples, were between 0.9999 and 1.000 and the concentra- tion RSD of the data were mostly less than 3%. The whole- Mineral chemistry of the basalts of the Iwatsubodani rock compositions of the samples are listed in Table 1, and Formation are displayed on variation diagrams for selected elements against SiO2 in Fig. 6. Clinopyroxene and plagioclase in basalts A and B were analyzed by the Electron Probe Micro Analyzer (EPMA; Basalt A (basalt with a sub-ophitic texture) JEOL JXA-8800) at Nagoya University, using natural and synthetic standard minerals and a conventional procedure. The SiO2 concentration is between 48 and 51 wt%, and The minerals are generally homogenous in a single grain, deviations of each element are <1.3 wt% for major ele- and the core of each grain was analyzed. The chemical ments, <24 ppm for trace elements except for Sr, and composition of the clinopyroxene and plagioclase in basalts <1.1 ppm for REE. The loss on ignition (LOI) of the sam- A and B is shown in Table 2. ples is less than 2 wt%. The samples are nearly identical to each other in their concentrations of the following ele- Mineral composition of basalt A ments: TiO2 ca. 1.3 wt%; Al2O3 ca.13 wt%; MnO ca. 0.22 wt%; MgO ca. 6.5 wt%; P205 ca. 0.11 wt%; Cr ca. 80 ppm; Clinopyroxene varies from 53 to 72 in Mg# (=100 Mg/ Co ca. 45 ppm; Zn ca. 81 ppm; Ga ca. 16 ppm; and Nb (Fe+Mg)), (average 62). Ca/(Ca+Fe+Mg) is between ca. 3.8 ppm (Fig. 6). The samples are slightly increased in 0.30 and 0.45, (average 0.38). The TiO2 content varies in Na2O, K2O, Sr, and Zr, and reduced in Fe2O3*, V, Y, and the range between 0.32 and 1.1 wt% (0.74 wt% average), Hf, with an increase in SiO2 (Fig. 6). Basalt A has a high while Cr2O3 is less than 0.17 wt%. Al203 ranges from 1.6 FeO*/MgO ratio ranging from 1.9 to 2.5, and is richer than to 4.4 wt%. Plagioclase is An54-67 (An = 100Ca/(Ca+ Na)). basalt B in TiO2, Fe2O3*, MnO, V, Co, Cu, Zn, Y, Nb, Hf, and Ta (Fig. 6; Table 1). Figure 7a shows a diagram of Mineral composition of basalt B MORB-normalized multi-element concentrations (called spidergram). The diagram shows a great enrichment in Rb The Mg# of clinopyroxene in basalt B ranges from 69 to and Ba, and a depletion in P compared with a representa- 85 (79 average) and is higher than that of basalt A. Ca/ tive MORB composition, though most of HFSE and HREE (Ca+Fe + Mg) is between 0.32 and 0.44, (average 0.40). form a fat trend which lies along the MORB. Nb and Ta The TiO2 content, which is from 0.29 to 0.64 wt% (0.39 do not exhibit any significant anomaly. The chondrite-nor- wt% average), is generally lower than that of basalt A, and malized REE patterns provide mostly fat characteristics shows a negative correlation with Mg#. AlzO3 content var- (Fig. 7b). ies from 2.2 to 2.8 wt% and Cr2O3 is less than 0.35 wt%. Basalt B (basalt with porphyritic texture) Discussion The SiO2 concentration is between 49 and 53 wt%, and deviations of each element are <2.2 wt% for major ele- Age of the Iwatsubodani Formation ments, <40 ppm for trace elements except for Cr and Sr, and <5.2 ppm for REE. The loss on ignition (LOI) of Nakama et al. (2010) reported a U-Pb age of ca. 280 Ma the samples is mostly less than ca. 4 wt%. TiO2 (ca. 0.60 from zircons in a gabbro/diorite of the mafic intrusive rock Springer 1436 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1.81 072204 Clast 0.319 4.98 072201 Clast 0.129 6.12 930704t Clast 8.28 930706c Clast 0.10 3.54 Clast Iw 6b Table 1Whlerock chemicalcmposition of thebasalts A andBdisplayed to thr sinifcant digits, excet forCZn,Ga, and Z 0.105 5.24 Clast Iw 6a 0.086 4.88 092003 092014 092709 Clast 0.0712 39 3 7.38 Clast 5. 0.12 5.57 Basalt B 15027216 81714210010440 1 0.090 5.57 IW 14 Lava 6851714681025152412 0.329 4.91 Lava Iw 9 0.270 4.35 Clast Iw 6c 40 65016 8161 72946383315040 0.300 5.05 9309101-10 Lava 44 18 5 10 10 4 6 3 0 4539 15 100 15 0.26 3.81 930804k Lava 4.06 Tajor element oxides (wt%) re-earth elements (ppm) 930804b Basalt A Lava 48.8 0.240 4.12 elements 1 Springer Int J Earth Sci(Geol Rundsch)(2017) 106:1429-1451 1437 072205 Clast 4.69 0.753 1A 072204 Clast 0.304 11.7 072201 Clast 14.0 930704t Clast 22.6 930706c Clast 10.1 Clast Iw 6b 17.1 3 0.188 Clast Iw 6a 2.63 C 16.0 13.9 601 0.172 092003092014092709 Clast 18.7 2.78 3 3.30 0.178 Clast #. 14.3 Basalt B Clast 12.2 2.3 0.960 IW 14 Lava 42 Lava 6MI 11.2 Clast Iw 6c S. 13.1 2. 9309101-10 Lava 0.629 4.26 0.927 10.1 0.367 930804b 930804k Lava 11.0 6.518 三 4.64 966°0 0.393 aFeO3, total ion as Fe2O3 Basalt A Lava Table 1 (continued) 10.9 0.396 0 m Springer 1438 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 (wt.%) (wt.%) (wt.%) 1.0 16- 口 12- 12 口 口P 0.8 口口 口 口 0.4 TiO2 AlO3 Fe203* ●basalt A > 14 . 12- 10 - 8 64 0.1 MnO MgO CaO 0 口 口 1.2- K20 0.2 Medium Potassium Series 0.15- 8 ●口 0.8- 口 口 中 Na20 0.4 P205 口 Low Potassium 0.05- 口 49 52 49 50 48 50 51 53 48 51 52 53 47 48 49 50 51 52 Si02 (wt.%) SiO2 (wt.%) SiO2 (wt.%) (ppm) (udd) (ppm) 500 口 口 ·日 300 Cr 400 250- 口口 40 200- 口 L 口 L 30 150 200 20- 100 100 Co 50- 10 250 100 20 200 16 ·品 口口 8 60- 口 150 口 口 口 口 12 100 40- Zn Cu 口 Ga 50 口 0 20 Rb Sr 口 300 口 200 口 口 100 ●basalt A √ basalt B 48 49 50 52 53 49 50 51 52 47 49 51 Si02 (wt.%) SiO2(wt.%) SiO2 (wt.%) Fig. 6 Variation diagrams of selected elements Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1439 (ppm) (udd) 100 80 Zr 口 60 口 口 40 口 Nb 20 Pb ●basalt A basalt B Ba 150 口 100 0.2 口 口 口 口 口 口 口 Hf 口 0.1 Ta 口 $0 口· 日 口 47 48 49 50 51 52 53 47 48 49 50 51 52 53 47 48 49 50 51 52 53 Si02 (wt.%) SiO2 (wt.%) SiO2 (wt.%) Fig. 6 (continued) b 70 50 basalt A basalt A Z 10 70 70 50 basalt B basalt B C1-chondri 70 mean mean basalt A basaltA basaltB basalt B 0.1 Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Fig. 7 a MORB-normalized multi-element concentration diagrams ough (1989) is used. b Chondrite-normalized REE patterns of the of the examined samples. The order of elements is after Pearce and examined samples. Normalizing chondrite concentrations are after Stern (2006). Normalizing MORB composition by Sun and McDon- Anders and Grevesse (1989) and Yamamoto et al. (2005) Springer 1440 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 Table 2 Chemical composition of plagioclase and clinopyroxene Occur- Basalt A rence Lava mineral P1 Pl P1 PI P1 P1 P1 Cpx Cpx Cpx Cpx Cpx Cpx Cpx SiO2 53.56 54.87 52.87 53.45 51.41 52.53 53.34 50.45 49.75 50.04 50.82 49.38 49.06 50.55 TiO2 b.d.1 b.d.1 b.d.1 b.d.1 b.d.1 b.d.1 b.d.1 0.47 0.70 0.75 0.44 1.04 0.95 0.66 Al2O3 28.13 27.70 28.35 27.85 29.56 28.88 28.24 3.48 2.37 1.79 1.88 2.73 2.71 2.80 Cr203 b.d.1 b.d.l b.d.1 b.d.l b.d.1 b.d.l b.d.1 0.10 b.d.l b.d.1 b.d.1 b.d.1 b.d.l b.d.1 FeOa 1.11 1.04 1.00 1.20 0.85 1.05 1.13 12.37 15.84 19.84 14.07 17.56 16.36 10.99 MnO b.d.l b.d.1 b.d.1 b.d.1 b.d.l b.d.l b.d.l 0.32 0.40 0.48 0.36 0.41 0.31 0.26 MgO 0.14 0.16 0.17 0.16 0.25 0.17 0.17 16.79 13.52 12.61 15.02 12.49 11.82 14.68 Cao 11.87 11.27 12.11 12.00 13.85 12.65 12.08 14.94 16.62 14.08 15.75 16.21 18.38 18.79 Na2O 4.79 5.32 4.73 4.79 3.71 4.23 4.62 0.22 0.24 0.16 0.22 0.24 0.31 0.20 Total 99.60 100.35 99.23 99.45 99.62 99.50 99.58 99.15 99.44 99.75 98.56 100.05 99.91 98.94 0= 8 8 8 8 8 8 8 6 6 6 6 6 6 6 !S 2.435 2.468 2.411 2.434 2.348 2.397 2.427 1.895 1.907 1.932 1.940 1.894 1.888 1.911 Ti 0.013 0.020 0.022 0.013 0.030 0.028 0.019 A1 1.507 1.468 1.523 1.494 1.591 1.553 1.514 0.154 0.107 0.082 0.085 0.124 0.123 0.125 Cr 0.003 / 一 一 一 Fe2+ 0042 0.039 0.038 0.046 0.033 0.040 0.043 0.388 0.508 0.640 0.449 0.563 0.526 0.347 Mn 0.010 0.013 0.016 0.012 0.013 0.010 0.008 一 一 Mg 0.010 0.011 0.012 0.011 0.017 0.012 0.012 0.940 0.772 0.726 0.855 0.714 0.678 0.827 Ca 0.578 0.543 0.591 0.585 0.678 0.618 0.589 0.601 0.683 0.582 0.644 0.666 0.758 0.761 Na 0.422 0.464 0.418 0.423 0.328 0.374 0.407 0.016 0.018 0.012 0.017 0.017 0.023 0.015 Total 4.994 4.993 4.993 4.993 4.994 4.994 4.993 4.021 4.028 4.011 4.013 4.022 4.034 4.014 Pl An 57.79 53.94 58.58 58.06 67.38 62.28 59.12 一 一 Mg# 70.76 60.34 53.12 65.55 55.91 56.30 70.43 Occur- Basalt A Basalt B rence Lava Clast mineral Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx SiO2 48.02 49.09 48.81 49.19 49.44 50.34 49.93 50.68 52.74 52.75 51.41 52.81 52.79 53.38 52.40 TiO2 0.32 1.12 0.96 0.59 0.81 0.82 0.20 0.64 0.43 0.29 0.43 0.30 0.38 0.30 0.33 Al2O3 4.07 4.80 4.24 2.68 4.43 3.85 1.63 2.76 2.21 2.53 2.85 2.31 2.27 2.24 2.25 Cr2O3 0.17 0.16 0.12 b.d.l 0.11 0.08 b.d.1 b.d.l 0.10 0.35 b.d.1 0.22 0.12 0.21 0.15 FeOa 14.58 11.70 12.57 14.01 17.15 10.92 14.81 12.85 9.27 6.58 9.60 7.60 8.50 6.61 9.43 MnO 0.46 0.36 0.35 0.50 0.32 0.20 0.48 0.30 0.27 0.18 0.33 0.21 0.26 0.19 0.31 MgO 13.40 14.96 13.36 12.49 12.56 15.40 9.93 16.29 16.35 16.42 15.23 16.84 16.11 16.29 16.21 CaO 15.53 16.81 18.25 17.87 13.06 17.81 21.14 15.20 19.25 21.27 19.81 19.65 19.82 21.14 19.30 Na2O 0.14 0.23 0.23 0.15 0.38 0.23 0.18 0.17 0.33 0.29 0.32 0.26 0.27 0.28 0.33 Total 96.67 99.23 98.87 97.47 98.25 99.58 98.30 98.89 100.85 100.29 99.96 99.96 100.40 100.42 100.56 0= 6 6 6 6 6 6 6 6 6 6 6 6 6 6 Si 1.880 1.850 1.863 1.915 1.904 1.883 1.952 1.913 1.936 1.930 1.915 1.940 1.941 1.949 1.932 Ti 0.009 0.032 0.028 0.017 0.023 0.023 900'0 0.018 0.012 0.008 0.012 0.008 0.010 0.008 0.009 Al 0.188 0.213 0.191 0.123 0.201 0.170 0.075 0.123 0.096 0.109 0.125 0.100 0.098 0.096 0.098 Cr 0.005 0.005 0.004 0.003 0.002 0.003 0.010 0.007 一 0.004 0.006 0.004 Fe2+ 0.477 0.369 0.401 0.456 0.552 0.342 0.484 0.406 0.238 0.161 0.233 0.209 0.239 0.197 0.230 Mn 0.015 0.011 0.011 0.016 0.010 0.006 0.016 0.010 0.008 0.006 0.010 0.006 0.008 0.006 0.010 Mg 0.782 0.841 0.760 0.725 0.721 0.859 0.579 0.916 0.894 0.896 0.846 0.922 0.883 0.887 0.891 Ca 0.651 0.679 0.746 0.745 0.539 0.714 0.885 0.615 0.757 0.834 0.791 0.773 0.781 0.827 0.762 Na 0.011 0.017 0.017 0.011 0.028 0.017 0.014 0.006 0.012 0.010 0.011 0.009 0.010 0.010 0.012 Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1441 Table 2 (continued) Occur- Basalt A Basalt B rence Lava Clast mineral Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Cpx Total 4.018 4.017 4.020 4.010 3.9814.016 4.011 4.006 3.956 3.963 3.943 3.974 3.975 3.987 3.947 PI An Mg# 62.10 69.51 65.46 61.39 56.6471.54 54.46 69.32 78.98 84.73 78.40 81.54 78.68 81.79 79.48 Pl plagioclase, Cpx clinopyroxene, An anorthite Mg# = 100 × Mg/(Mg+Fe) aFeO is total ion as FeO ("dolerite/gabbro” in their paper) along the Iwatsubodani Tectonic setting of the Iwatsubodani Formation Valley (Fig. 4) and, therefore, assigned a Permian age to the Iwatsubidani Formation. However, the dated gabbro/ The geochemistry of basalt gives evidence for the tectonic diorite locality is within the distribution area of the rocks of setting of the volcanic activity that formed it because the the Hitoegane Formation (Fig. 4), and therefore, the age of chemical composition of basalt varies according to its ori- the gabbro/diorite does not necessarily indicate the age of gins, and many discrimination diagrams of basaltic rocks volcanic activity of the Iwatsubodani Formation, because have been proposed (e.g., Pearce 1982). In this section, the relationship between the dated rock and the mafic vol- discrimination diagrams are used to discuss the tectonic canic rocks of the Iwatsubidani Formation is uncertain. In setting and magma type of the basalts of the Iwatsubodani addition, the fine-grained facies of the “gabbro/diorite" Formation. include idiomorphic phenocrysts of hornblende and plagio- Na, K, Ca, Rb, Ba, and Sr are known to be highly mobile clase more than 1 mm in major axis, so it is quite different through secondary processes. Discrimination of the ana- from the Iwatsubodani basalt in both texture and mineral lyzed basalts is mainly based on the diagrams using rela- composition (Fig. 5c, e, f). Thus, the suggestion that the tively fuid-immobile elements, such as Ti, Fe, Mg, P, V, Y, gabbro/diorite and the rocks of the Iwatsubodani Formation Zr, Nb, and REE. s the evidence (Tsukada 1997). Hence, it is not appropriate Tectonic setting of basalt A to link the age of the Iwatsubodani Formation to that of the gabbro/diorite from the stratigraphic or the petrographic Basalt is generally formed at a mid-oceanic ridge, a vol- viewpoint. Tsukada (1997) suggested that the Iwatsubodani canic arc, or in an intra-oceanic/-continental plate (i.e., Formation was primarily overlain by the Hitoegane Forma- within-plate). It can be divided into alkaline and non- tion based on the following facts: (1) the lower part of the alkaline basalts, the latter further subdivided into tholei- Hitoegane Formation includes many mafic volcanic clasts; itic and calc-alkaline basalts by their magma source (2) the Hitoegane Formation intercalates mafic pyroclastic variations and resulting chemical compositions (e.g., rock layers, including clasts of both A and B basalts; (3) Miyashiro and Kushiro 1975; Pearce 1982). The HFSE the sedimentary structures of both the Iwatsubodani and vs. La/Yb diagrams of basalt A (Nakamura et al. 2000) Hitoegane Formations indicate northwestward-up around suggest that the samples are non-alkaline basalt (Fig. 8). their boundary; and (4) the strike and dip of the bedding In some diagrams with relatively immobile elements, planes of the Iwatsubodani and Hitoegane Formations are such as Ti, P, Y, Zr, and Nb, the data are plotted in the mostly the same around their boundary (Fig. 4). Llanvir- field of MORB or volcanic arc basalt (VAB) (e.g., Mul- nian to Caradocian conodonts, Periodon aculeatus, were len 1983; Pearce 1982, 1983; Pearce and Cann 1973, obtained from the lower part of the Hitoegane Forma- Fig. 9). Both the Zr/Y vs. Zr diagram (Pearce 1983) and tion (Tsukada and Koike 1997), and Nakama et al. (2010) theTi vs. Zr diagram (Pearce 1982） indicate MORB reported zircon ages of ca. 472 Ma, assigned from the latest or VAB (Fig. 9a, b). In addition, the data cluster in the Cambrian/earliest Ordovician to Caradocian (International feld of low-potassium tholeite (LPT) of volcanic arc Commission on Stratigraphy ed. 2015), from felsic tuff of in the Zr-Ti/100-3Y diagram (Pearce and Cann 1973; the lowest horizon of this formation (Fig. 4). These lines of Fig. 9c). The LPT and CAB in this diagram do not nec- evidence all point to the age of the Iwatsubodani Formation essarily mean that the rock originated from magmas of being earlier than late Ordovician. tholeite and calc-alkaline series, because the results Springer 1442 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 ppm ppm 10 10 TiO2 Hf 0.1 0.1 ppm ppm 1000 10 Zr Ta 100 0.1 10 0.01 0.1 10 100 ppm La/Yb 1000 MORB Volcanic arc basalt Nb 100 Island arc tholeite Island arc alkaline Calc-alkaline 10 Back arc basin Within-plate basalt Oceanic island tholeite Oceanic island alkaline 0.1 0.1 10 > Intra-continental tholeite 1 100 La/Yb ; Intra-continental alkaline Fig. 8 HFSE vs. La/Yb diagrams (Nakamura et al. 2000) are very similar to those from the magmas of low- and Fig. 9d), and the MnO-TiO2-P2O5 diagram (Mullen medium-potassium series basalt in volcanic arcs (e.g., 1983) also indicates their island arc tholeite (IAT) affin- Ujike 1989). Taking all these lines of evidence together, ity (Fig. 9e). it is reasonable to conclude that the rocks were formed One of the important geochemical indicators in volcanic in low-potassium series magma at a volcanic arc. In the arc basalts is depletion in HFSE (esp. Nb and Ta) com- SiO2 vs. FeO*/MgO diagram, SiO2 of basalt A is almost pared with LILE and LREE, a feature that is not observed kept constant regardless of their FeO*/MgO ratio, sug- in MORB (e.g., Gill 1981; Pearce et al. 2005). The spider- gesting tholeiitic affinity of the magma (Miyashiro 1974; gram for basalt A samples shows an enrichment of LILE Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1443 Fig. 9 Discrimination diagrams ●basalt A of a Zr/Y vs. Zr (Pearce 1982), a 10 Ti/100 VAB basalt B b Ti vs. Zr (Pearce 1983), c Zr- WPB Ti/100-3Y (Pearce and Cann 人/Z 1973), d SiO2 vs. FeO*/MgO (Miyashiro 1974), and e MnO- % MORB TiO2-P2O5 (Mullen 1983) WPB 100 150 (ppm) LPT 10 Zr b N-MORB (ppm) LPT CAB CAB WPB Zr 3Y 10000 Ti 占 VAB e TiO2 1000 10 100 1000 Zr (udd) (wt.%) 60 MORB d 口 CA OIT 55 TH SiO2 IAT OIA 50 L CAB 45 0.5 1.5 2.5 10MnO 10P2O5 1 FeO*/MgO in comparison with HFSE, similar to volcanic arc basalt. 1.5 However, the HFSE and HREE of basalt A show a flat the Ti, Cr, and Ca in clinopyroxene phenocrysts suggest trend as seen in the MORB, and there is no negative Nb 0.5 and Ta anomaly (Fig. 7a). A, while on the other hand, the clinopyroxene Mg# vs. pla- In chondrite-normalized REE patterns, MORB and gioclase An diagram give a clear indication that it is IAT IAT generally show a horizontal fat line with a compara- rather than MORB. How can discrimination diagrams for tively wide range in LREE, while the CAB shows a grad- whole-rock composition and clinopyroxene Mg# vs. pla- ual HREE depletion (Nakamura et al. 2000 and references s therein). The present samples show horizontal flat REE basalt, whereas HFSE and REE show MORB-like nature patterns which are similar to MORB or IAT (Fig. 7b). Ti in spidergram (Figs. 7a, 9, 10)? Lack of a negative Nb vs. (Ca+ Na) and (Ti+ Cr) vs. Ca diagrams for clinopyrox- anomaly in the spidergram is also a significant characteris- tic of MORB-like rocks (Fig. 7a). Several authors reported genic tholeiite”’ of which the most important type is MORB "MORB-like tholeitic fore-arc basalt (FAB)" from, for (Fig. 10a, b). It has long been known that the composition example, the Izu-Bonin-Mariana (IBM) arc, which closely of major constituent minerals in basaltic rocks crystallized ppI from MORB magmas is clearly different from the composi- form FeO*/MgO ratio (average 2.2); moderate TiO2 (aver- 2.0 age 1.3 wt%); low LREE/HREE ratio (average 1.0); nearly Ishiwatari et al. 1990, 2003). For example, the clinopyrox- identical HFSE and HREE concentrations to those of ene is more magnesian in an MORB series in comparison MORB; a fat chondrite-normalized REE pattern; and the with an IAT series, at the same anorthite (An) content of 1.0 the coexisting plagioclase (Ishiwatari et al. 2 2006).The 80 clinopyroxene in basalt A has a lower Mg# than that from Ishizuka et al. 2009, 2011, 2014; Li et al. 2013; Reagan normal MORB, and clinopyroxene-plagioclase pairs of the et al. 2010; Table 1; Fig. 7). It has been recognized that V Springer 1444 Int J Earth Sci (Geol Rundsch) (2017) 106:14291451 Fig. 10 Diagrams for clinopy- 0.15 0.06 roxene and plagioclase chemical a 0.05 HL oquso1o-uou b compositions. a, b Diagrams for clinopyroxene phenocryst (Leterrier et al. 1982). The sam- basalt A 0.03 ples plotted in the field of “TH and CA basalt” are adaptive to basalt B T the diagram (b). TH tholeite, 0.02 品 orogenic CA calc-alkaline. c Clinopyrox- 0.1 basalt 10'0 ene Mg# (=100 Mg/(Fe + Mg)) and plagioclase An value (=100 alkaline basalt 0 Ca/(Na + Ca)) coexisting in the 0.5 0.6 0.7 Ca 0.8 0.9 !L 1.0 sample from basalt A. Nearly identical data were obtained TH and CA C from three samples of the basalt basalt 90 A from the pillow lava, and 0.05 their average is shown here. MORB 80 Samples obviously have an affinity to the island arc tholeite 70 IAT (IAT) rather than normal MORB. MORB and IAT fields 60 ★ are after Ishiwatari (1999) and Clin basalt A Ishiwatari et al. (2006) 50 0.5 1.0 50 60 70 80 90 100 + Plagioclase average An concentration and Ti/V ratio are significant factors to dis- subduction-derived fuids (Beccaluva et al.2005;Reagan criminate between FAB and typical IAT and MORB (e.g., et al. 2010) and the inference is that this may also be the Ishizuka et al. 2009; Reagan et al. 2010). FAB is known case for basalt A. to have a comparativelyhigh V concentration (generally more than 200 ppm) and to show a lower Ti/V ratio (com- Tectonicsettingof basalt B monly 1020) than MORB (e.g., Ishizuka et al. 2009; Rea- gan et al. 2010). In addition, it generally shows a higher The HFSE vs. La/Yb diagrams suggest a high possibility Ti/Vratio than IAT when V concentration is higher than of calc-alkaline basalt (CAB) for the samples (Nakamura ca. 350 ppm (e.g. Beccaluva et al. 2005; Ishizuka et al. et al. 2000; Fig. 8). Basalt B is plotted in the fields of the 2011,2014;Reagan et al. 2010;Fig. 11).In the present MORB or LPT or CAB in the Ti-Zr-Y diagram Pearce samples, a V concentration of 364456 ppm and a Ti/V and Cann 1973; Fig.9c).In other words, this basalt is ratio of between 18 and 21 suggest an affinity to FAB but characteristic of MORB or arc basalt from magma of not to normal MORB or IAT (Fig. 11). When the basalt A low-to-medium-potassium series.Discrimination dia- samples are plotted on the V vs. Ti/1000 diagram (Fig. 11), grams using Ti, Zr, and Y tell us that basalt B is volcanic they overlap with the field of FAB and back-arc basin basalt arc basalt (Pearce 1982,1983; Fig.9a,b). The trend in (BABB). However, most BABB has a Ti/V ratio of more the spidergram,showing an enrichment in LILE com- than 20 (Shervais 1982), which suggests a higher probabil- pared with HFSE, supports the results of the Ti, Zr, and Y ity of basalt A being FAB rather than BABB.In general. examination (Pearce 1983;Fig. 7a). A pronounced nega- FAB has a lower LREE/HREE ratio in chondrite-normal- tive Nb and Ta anomaly in the spidergram also suggests ized values than MORB or BABB from the Philippine Sea volcanic arc basalt (Fig. 7a).The chondrite-normalized Basin (Hickey-Vargas 1998;Ishizuka et al. 2009, 2014), REE patterns with a gradual HREE depletion are similar and it lacks a clear enrichment in LILE compared to REE to those of CAB (Nakamura et al. 2000 and references and HFSE, implying it received little or no input from slab- therein,Fig.7b).The SiO vs.FeO*/MgO (Miyashiro derived material (Ishizuka et al. 2014). However, basalt A 1974)and MnO-TiO-P2Os(Mullen 1983)diagrams differs from this in showing a slightly higher LREE/HREE indicate CAB for basalt B (Fig. 9d, e). This basalt is ratio than MORB,and in being enriched in LILE (e.g.,Rb similar to calc-alkaline rocks,such as those of the Bonin and Ba) compared with HFSE and REE (Table 1; Fig. 7a). Islands, in terms of its distinctive LILE-enrichment, Nb LILE- and/or LREE-enriched FAB, such as those from the depletion,moderate V concentration,and low Ti/V ratio DSDP sites 458 and 459 and from the AlbanideHellenide (Ishizuka et al. 2014; Shervais 1982;Figs. 7a, 11). The orogenic belt in the Balkans,is considered to have been Ti vs.(Ca+Na) diagram for clinopyroxene phenocryst formed as a result of the modification of the melt source by Springer Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 1445 ppm Petrologic-tectonic variation of the Hitoegane 500 0 20 succession Mariana- IAT Bonin FAB The chondrite-normalized REE patterns clearly demon- strate that basalt B has a much higher LREE/HREE ratio 400 than basalt A and that basalt A has higher HREE concen- MORB trations than basalt B (Fig. 7b). These facts suggest that basalts A and B were formed by different magmatic pro- grams seem to show that the FAB volcanism that gave 300 rise to basalt A and the calc-alkaline volcanism resulting BABB V in basalt B took place in a volcanic arc. With this back- Alkalinebasalt ground, the environmental variation of the Hitoegane suc- and OIB CAB cession will now be discussed. In the Iwatsubodani Forma- 200 tion, pyroclastic rock that includes abundant clasts of basalt B with minor amounts of basalt A overlies pillow lava entirely composed of basalt (A) The chemical composition of basalt B, which forms the majority of the clasts in the 100 pyroclastic rock, reflects the volcanism of the pyroclastic ●basalt A rock, and the composition of basalt A, which forms the pil- TH and CA arc volcanic rocks from basalt B low lava, indicates a slightly earlier volcanism than the vol- Bonin Islands canism of basalt (B). In other words, the early FAB volcan- ism that gave rise to basalt A was followed by the high-Mg 5 10 15 calc-alkaline volcanism of basalt B. Ti/1000 The stratigraphic succession composed of FAB, bonin- ite, and high-Mg andesite, and tholeitic/calc-alkaline vol- Fig.11 V vs. Ti/1000 diagram (Shervais 1982). The IAT, CAB, canic rocks in ascending order has been reported world- BABB, MORB, Alkaline basalt, and OIB fields are after Shervais (1982), and others are after Ishizuka et al. (2014). The diagram sug- wide from supra-subduction zones (SSZ) (e.g., Beccaluva gests that the basalt A is fore-arc basalt (FAB) rather than normal et al. 2005; Dilek and Furnes 2003; Ishizuka et al. 2014). MORB. Although the basalt A is plotted in the overlapped field of the The Iwatsubodani Formation with FAB and overlying FAB and back-arc basin basalt (BABB), the majority of the BABB (85% data within Shervais 1982) is more than 20 in Ti/V ratio; there- high-Mg CAB is quite similar to this standard succession, fore, the basalt A is more probable to FAB than BABB. IAT typical 0.4 island arc tholeite, OIB oceanic island basalt, CAB calc-alkaline however, that high-Mg andesite (58 wt% SiO2, 0.90 wt% basalt, TH tholeite, CA calc-alkaline + itic rock (53 wt% SiO2, 0.49 wt% TiO2, 8.5 wt% MgO and (Leterrier et al. 1982) clearly indicates that basalt B is 1.1 FeO*/MgO ratio) were obtained from river boulders in non-alkaline. The (Ti+Cr) vs. Ca diagram (Leterrier the upper reaches of the Iwatsubodani Valley (Fig. 4; Tsu- et al. 1982) suggests “orogenic basalt'", including island kada, unpublished data). It follows that high-Mg andesite arc tholeiitic, calc-alkaline, and shoshonitic basalts and boninite are likely to occur in the Iwatsubodani For- mation, but this remains to be confirmed, because the geo- a result of an examination of both whole-rock and clino- logical relationship between the “High-Mg andesite and pyroxene chemistries. Most basalt B samples, 092014, boninite”’ and this formation is still uncertain. Based on 092709,Iw 6a,Iw 6b,930706c,072201,and 072204, lithological, chronological, and geochemical examination, resemble the high-Mg basalt of the Miocene Northeast the model proposed is that FAB is erupted at the earliest Japan arc in their high MgO (average 8.5 wt%), high Cr stage of island arc formation (the initial stages of subduc- (average 193 ppm), and low FeO*/MgO ratio (average tion) and that subsequently boninitic/tholeitic/calc-alka- 1.2) (Shuto et al. 1985, 2015; Takimoto and Shuto 1994; line arc volcanism occur at an SSZ (e.g., Dilek and Furnes Table 1), and the latter two samples seem to be ankaram- 2014; Ishiwatari et al. 2006; Ishizuka et al. 2014; Li et al. ite in high CaO concentration. Sample 930706c is picritic 2013). If this model is applied to the Iwatsubodani For- with 11.6 wt% in MgO, 238 ppm in Cr, and 149 ppm in 0.8 Ni. The conclusion is that the basalt B is of CAB, and Paleozoic subduction initiation volcanism and the overly- some samples are probably high-Mg CAB. 25 volcanism at an SSZ. Springer 1446 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 The Iwatsubodani Formation is entirely composed primitive oceanic arc origin for this ophiolite (e.g., Arai of mafic volcanic rocks, but in contrast, the overlying and Yurimoto 1995; Ishiwatari 1989; Tsujimori and Itaya Hitoegane Formation includes large amounts of felsic tuff 1999). Ehiro et al. (2016) proposed that this ophiolite had and tuffaceous clastic rocks yielding Periodon conodonts, acted as a fore-arc mantle wedge and as a basement for lower Paleozoic fore-arc basin formations. Jadeitite within slope environment (e.g., Armstrong et al. 2001; Dubinina the ophiolite showing high-pressure and low-temperature and Ryazantsev 2008; Tsukada 1997; Tsukada and Koike metamorphism probably suggests a fore-arc mantle wedge 1996; Fig. 4). The lower part of the Hitoegane Formation at an SSZ (Ehiro et al. 2016; Stern et al. 2013; Harlow et al. intercalates mafic pyroclastic rock layers similar to those of 2015). In either scenario, it is now accepted that the Oey- the Iwatsubodani Formation (Tsukada 1997), and this sug- ama ophiolite is of volcanic arc origin at an SSZ. Gabbroic gests a conformable/unconformable relationship between intrusive rock in the ophiolite gives Sm-Nd isochron ages the Iwatsubodani and Hitoegane Formations, as previ- of ca. 560 Ma (Hayasaka et al. 1995), and jadeitite yields ously discussed. The zircon U-Pb age of ca. 472 Ma from hydrothermal zircon U-Pb ages ranging from ca. 520 to the lowest horizon of the Hitoegane Formation suggests 450 Ma (Ehiro et al. 2016; Tsujimori et al. 2005; Kunugiza that the volcanic transition from mafic to felsic was likely and Goto 2010). In addition, peridotite and amphibolite/ to have occurred around the early or middle Ordovician metagabbro block in the peridotite give hornblende K-Ar (International Commission on Stratigraphy 2015; Nakama ages of ca. 464-444 Ma, and ca. 443-403 Ma, respectively et al. 2010). The upper part of the Hitoegane Formation (Nishimura and Shibata 1989; Tsujimori 1999; Tsujimori yields upper Silurian radiolarians (Pseudospongoprunum et al. 2000; Tsujimori and Ishiwatari 2002). Ishiwatari and tauversi and Futobari solidus-Zadrappolus tenuis assem- Tsujimori (2003) regarded this ophiolite as Cambro-Ordo- blages) and contemporaneous zircons of ca. 426 Ma (Man- vician. In summary, the Oeyama ophiolite, located between chuk et al. 2013b). The upper Silurian tuffaceous rocks the Hida and Mino-Tamba belts, is considered to have been of the Fukuji succession in the Fukuji area, 3 km west of formed at a volcanic arc at an SSZ in the early Paleozoic, the study area (Fig. 3), yields radiolarians of the Futobari as is the case with the Iwatsubodani Formation of the Hida solidus-Zadrappolus tenuis assemblage and contempora- Gaien belt. Sugamori and Ishiwatari (2015) proposed a neous zircons of ca. 421 Ma (Manchuk et al. 2013a). This model in which the Permian marginal sea (called the Mai- lithological, paleontological, and chronological evidence zuru Ocean) was formed as a result of back-arc-spreading strongly suggests that the upper Silurian tuffaceous rocks which split the older arc crust that constituted the “proto- of the Fukuji succession are correlated with a part of the Oeyama ophiolite.’ Further they speculated that the Oey- Hitoegane Formation (Manchuk et al. 2013b; Fig. 2). ama ophiolite and the back-arc spreading-related Permian ophiolite (the present Yakuno ophiolite in the Maizuru belt, Regional correlation of the Hitoegane succession e.g., Ichiyama and Ishiwatari 2004), are the “fossilized remains"’ of the Paleozoic arc-marginal sea system which The previous section proposed that the Hitoegane succes- had existed for 300 million years. This Paleozoic arc-mar- sion, composed mainly of mafic volcanic rock (Iwatsub- ginal sea system might have been generated by the subduc- odani Formation) and clastic and felsic tuffaceous rocks tion of an oceanic plate which was initiated in the early (Hitoegane Formation), was formed at an early Paleozoic Paleozoic as evidenced by the basaltic rocks reported here. volcanic arc at an SSZ. Here, the regional correlation of the The lithostratigraphic similarity between the Hitoegane Hitoegane succession will be examined with a key word, succession and the Hayachine-Miyamori ophiolite and i.e., an early Paleozoic volcanic arc at an SSZ. covering formations (called the H-M succession, here) of The lower Paleozoic Oeyama ophiolite that structurally the South Kitakami belt, northeast Japan, has been pointed overlies the rocks of the Renge metamorphic rocks and out (e.g., Ehiro et al. 2016; Tazawa 1993; Tsukada and Akiyoshi, Maizuru, and Mino-Tamba belts is sporadically Koike 1997; Fig. 1). The Hayachine-Miyamori ophiolite exposed in the Sangun-Renge belt, Chugoku Mountains, is composed mainly of peridotite, pyroxenite, hornblendite, west of Hida Gaien belt, southwest Japan (Fig. 1). The gabbro, and mafic volcanic rocks (e.g., Ehiro et al. 2016; Oeyama ophiolite is composed mainly of residual perido- Ozawa et al. 2015). The Hayachine-Miyamori ophiolite has tite with podiform chromite deposits and minor gabbroic a common petrologic feature, namely, the presence of horn- rocks, while mafic volcanic rocks are not yet found (e.g., blende as a major constituent mineral in both mafic and Ishiwatari and Tsujimori 2003). Amphibolite and jadei- ultramafic rocks, which suggests a hydrous condition for the tite occur as blocks or dykes in places (e.g., Ehiro et al. ophiolite formation (Ozawa 1984, 1988). At one time, this 2016; Kurokawa 1985; Nishimura and Shibata 1989; Tsu- was inferred to have been formed in a rift zone (e.g., Osawa jimori and Liou 2004). The chemical composition of spi- 1983), however, it was pointed out that the geochemical and nel in peridotite and chromite suggests back-arc basin or petrological features suggest an arc environment rather than Springer Int JEarth Sci(GeolRundsch)(2017)106:1429-1451 1447 MORB equivalents (e.g., Mori et al. 1992; Ozawa 1984). clastic and felsic tuffaceous rocks, and similarly, the Ozawa et al. (2015) proposed that the Hayachine-Miy- Hitoegane succession is composed of basalt (Iwatsub- amori ophiolite is divided into the back-arc “Hayachine odani Formation) and the overlying Ordovician-Silurian complex" and the fore-arc “Miyamori complex" Ozawa clastic and felsic tuffaceous rocks (Hitoegane Formation) et al. (2015) considered that the aluminous spinel ultra- (e.g., Tazawa 1993; Tsukada and Koike 1997). mafic suite (ASUS) in the Hayachine complex underwent Accordingly, the H-M succession and the Hitoegane melting in the back-arc region and that the ASUS and succession are quite likely correlated in the similarity chromite-bearing ultramafic suite of the Miyamori com- of their ages (Early Paleozoic), their tectonic settings plex underwent melting closer to the fore-arc region. In (arc volcanism at SSZ), and their environmental varia- the opinion of Ozawa et al. (2015), however, these com- tions (mafic volcanism and following felsic volcanism). plexes are not two separate ophiolites that are genetically Okawa et al. (2013) reported Precambrian detrital zircons unrelated to each other, but rather that they share the same from the Yakushigawa Formation and its equivalent, and arc-back-arc system in an SSZ. Such a back-arc origin for assumed that these formations were formed near a conti- the Hayachine complex is consistent with the conclusions nental magmatic arc having Precambrian basement rocks, of Ehiro and Kanisawa (1999) and Uchino and Kawamura at around northern East Gondwana. (2016). Shibata and Ozawa (1992) reported a 510± 70 Ma A similarity between the Oeyama ophiolite, the Hay- Sm-Nd whole-rock isochron age for the Miyamori com- achine-Miyamori ophiolite, and the Sergeevka ophi- plex, and Yoshikawa and Ozawa (2007) presented a clino- olite (Promorye, Far East Russia) has been recognized pyroxene Sm-Nd model age of 499 ±65 Ma. Yoshikawa in many aspects (e.g., Ishiwatari and Tsujimori 2003; et al. (2012) obtained a clinopyroxene Sm-Nd isochron of Fig. 1). The Sergeevka ophiolite structurally overlies 490 ± 60 Ma from the Hayachine and Miyamori complexes, the younger blueschist-facies metamorphic rocks (Shai- and Ozawa et al. (1988) and Shibata and Ozawa (1992) ginskiy blueschist) and Jurassic-Cretaceous accretionry estimated, based on hornblende K-Ar method, the cooling complexes (Samarka terrane). The Shaiginskiy blueschist age of this ophiolite as 400-500 Ma. Zircon LA-ICP-MS and Samarka terrane correspond to the Renge metamor- U-Pb ages of 466±6 Ma and 457±10 Ma were obtained phic rocks and the Mino-Tamba belt of Southwest Japan, from the trondhjemite and the felsic tuff of the Hayachine respectively, and the Sergeevka and Oeyama ophiolites complex (Shimojo et al. 2010). Taking all these together, share the same structural position in that they both over- the ages by hornblende K-Ar, hornblende Sm-Nd, clino- lie the blueschist-facies metamorphic rocks and Mesozoic pyroxene Sm-Nd, and Zircon U-Pb methods, therefore, accretionry complexes (e.g., Ishiwatari and Tsujimori give a concentration at 400-500 Ma (Ozawa et al. 2015). It 2003; Kojima 1989; Kojima et al. 2000). The Sergeevka has generally been considered that the Ordovician to Silu- ophiolite is composed mainly of gneissose hornblende rian clastic rocks and felsic tuffaceous rocks (Yakushigawa - - s 1 Formation） conformably cover the Hayachine complex (Khanchuk et al. 1996). Mishkin et al. (1970) reported (e.g., Ehiro et al. 2016). However, Shimojo et al. (2010) muscovite with a K-Ar age of 529 Ma and hornblende proposed an unconformable relationship between the Hay- K-Ar age of 622 Ma, whereas Ishiwatari and Tsujimori achine complex and the Yakushigawa Formation on the (2003) presented hornblende K-Ar ages of 470-430 Ma. grounds that detrital zircons of ca. 425 Ma are contained in The conclusion is that the Segeevka, Oeyama, and Haya- the lower part of this formation. Upper rocks (the Odagoe chine-Miyamori ophiolites are correlated with each other Formation and its equivalent) of the Yakushigawa Forma- in that they have a similar structural position, lithology, tion yield Silurian coral, trilobites, and brachiopods (e.g., and age (Ishiwatari and Tsujimori 2003). Ehiro et al. 1986, 2016; Okami et al. 1986). Ozawa et al. As has been mentioned, the Iwatsubodani Formation of (2015) concluded that the Hayachine-Miyamori ophiolite the Hida Gaien belt and the Oeyama ophiolite is probably was formed at an Ordovician SSZ arc-back-arc system and comparable to each other. Thus, if the Oeyama ophiolite is that subsequently Ordovician (?)-Silurian felsic volcanism a relative of the Sergeevka ophiolite, the Iwatsubodani For- occurred at this area as evidenced by the Yakushigawa and mation also might be the equivalent of the Sergeevka body. Odagoe Formations. Besides, the Hitoegane succession corresponds to the H-M The Hayachine-Miyamori ophiolite was formed at succession in environmental variation: early Paleozoic SSZ early Paleozoic volcanic arc at an SSZ as was the case mafic arc volcanism followed by felsic volcanism. Conse- for the Iwatsubodani Formation, as mentioned above. In quently, the evidence points to the fact that the Hitoegane addition, it has been pointed out that there is a strong succession of the Hida Gaien belt, the Oeyama ophiolite, stratigraphic similarity between the H-M succession the H-M succession, and the Sergeevka ophiolite probably and the Hitoegane succession: the Hayachine-Miyamori shared an oceanic arc system at an SSZ near a continental ophiolite is overlain by the Ordovician (?) or Silurian margin in the early Paleozoic. Springer 1448 Int J Earth Sci (Geol Rundsch) (2017) 106:1429-1451 Acknowledgements We would like to express our thanks to the late Faure M, Caridroit M, Charvet J (1986) The Late Jurassic oblique col- These similarities make it difficult to lisional orogen of SW Japan: new structural data and synthesis. University of Toyama for their critical advice. We would also like to Tectonics 5:1089-1114 relatively thin oceanic-type crust rather than Gill JB (1981) Orogenic andesites and plate tectonics. Springer, p 390 Prof. M. Takeuchi, and Prof. H. Yoshida at Nagoya University for and lead such authors as JAKEs & GILL helpful discussion. We appreciate critical review comments from Dr. tonics. Ann Rev Earth Planet Sci 43:105-138 A. Ishiwatari of the Nuclear Regulation Authority of Japan and an Hayasaka Y, Sugimoto T, Kano T (1995) Ophiolitic complex and anonymous reviewer, which greatly improved an earlier draft of this metamorphic rocks in the Nimi-Katsuyama area, Okayama manuscript. 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Ge0chem J 39:289-297 Tsujimori T (1999) Petrogenesis of the Fuko Pass high-pressure Yoshikawa M, Ozawa K (2007) Rb-Sr and Sm-Nd isotopic system- metacumulate from the Oeyama peridotite body, southwestern atics of the Hayachine-Miyamori ophiolitic complex: Melt Springer Int JEarth Sci(GeolRundsch)(2017)106:1429-1451 1451 generation process in the mantle wedge beneath an Ordovician Hayachine-Miyamori ophiolite: evidence for melting under island arc. Gondwana Res 11:234-246 influx of carbonate-rich silicate melt in an infant arc environ- Yoshikawa M, Suzuki K, Shibata T, Ozawa K (2012) Geochemical ment. J Min Petrol Sci 107:250-255 and Os isotopic characteristics of a fresh harzburgite in the Springer
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OKAYAMA UNIVERSITY Earth Science Report, Vol. 1, No. 1, 1-8, (1994) Petrography of primary peridotites from the Ohsa-yama area, Okayama Prefecture Toshio NozAKA and Tsugio SHIBATA Department of Earth Sciences, Faculty of Science Okayama University Ultramafic rocks exposed around Mt. Ohsa (= Ohsa-yama), Okayama Prefecture, designated as "Ohsa- yama ultramafic body" all together, are one of the Alpine-type peridotites in the Sangun metamorphic belt. They tion of the Ohsa-yama body where it has not been affected by the contact metamorphism, the constituent minerals, texture and structure of primary ultramafic rocks have been locally preserved. Petrographic studies reveal that the primary ultramafic rocks of the Ohsa-yama body consist dominantly of dunite and harzburgite possessing no obvious layering, and their constituent minerals are similar in composition to those of the Tari-Misaka and Ashidachi ultramafic bodies. These features indicate that unlike the Ochiai-Hokubo body, the Ohsa-yama ultra- mafic body belongs to the "massive group" of the Arai's (1980) classification. Keywords : petrography, dunite, harzburgite, massive ultramafic body 1. Introduction structure, and consist of dunite, harzburgite and A number of ultramafic bodies occur in the chromitite. He contends that the former repre- so-called "Sangun metamorphic belt" of high- sents cumulates derived from crystallization of pressure type in the inner zone of southwestern basaltic magma, the latter refractory residues Japan. Most of them are intensely serpenti- left after partial fusion of primordial peridotites, nized; in addition, they are in many cases ther- and both have initially constituted a part of an mally metamorphosed by younger granitic rocks ophiolite suite. (Research Group of Peridotite Intrusion, 1967). Compared to other regions in the Sangun Due to intense serpentinization and contact belt, ultramafic rocks are exposed most abun- metamorphism, the primary petrologic charac- dantly in the region from northwestern ters of these ultramafic bodies were generally ob. Okayama Prefecture to northeastern Hiroshima scured, but original structure, texture and min- Prefecture (Fig. 1). Among the relatively large erals have been preserved in some portions of in- ultramafic bodies in this region, the Tari-Misaka dividual bodies, hence making it possible to ob- and Ashidachi bodies belong to the Arai's mas- tain their primary characters (e.g., Igi and Abe, sive group, and the Ochiai-Hokubo body (also 1969; Arai, 1980). called "Onoro-yama body") belongs to the layered According to Arai (1980), the ultramafic group (Arai, 1980; Hamada, 1982); however, the bodies in the Sangun-Yamaguchi Zone can be nature of ultramafic rocks exposed around Mt. classified into two groups on the basis of original Ohsa (= Ohsa-yama*), Okayama Prefecture (des- structure and rock type; i.e., layered group and ignated hereafter as "Ohsa-yama ultramafic massive group. The layered bodies consist of sev- body") has not been clearly known. Petrography eral types of rocks such as lherzolite, dunite, of some peridotites in the Ohsa-yama body is pre- wehrlite, chromitite, websterite and clinopyroxe-: sented here for the purpose of clarification of its nite. The massive bodies rarely show layered primary petrologic nature. Toshio NozAKA and Tsugio SHIBATA 2 Pref 60N N TM 10 km 133°30E Fig. 1 Distribution of ultramafic rocks in the area of northwestern Okayama and northeastern Hiroshima Prefectures (compiled after Hiroshima Prefecture, 1964; Igi and Sakamoto, 1977; and Mitsuno and Sugita, 1980). Abbreviations for ultramafic bodies: TM, Tari-Msiaka body; AS, Ashidach body, OH, Ochiai-Hokubo body; OS, Ohsa-yama body. H1. Geological Setting is ca. 3 km wide from north to south and 5 km The Sangun metamorphic belt is charac- long from east to west. The exposed body has an terized by glaucophanitic metamorphism (e.g., unusual, ax-like shape in plane (Fig. 2). Rock Hashimoto, 1968). Recent chronological studies outcrops are found sporadically throughout the reveal that the Sangun metamorphic rocks have body. The ultramafic rocks are intensely serpen- a significantly wide range of radiometric ages tinized, except for those occurring at the south- (Nishimura and Shibata, 1987). On the basis of western slope of Mt. Ohsa where the serpenti- distribution of crystalline schists that show ap- nites have reverted to compact peridotitic rocks proximately identical ages, Nishimura (1990) due to contact metamorphism by a younger gra. subdivided the glaucophanitic terrane into three nitic intrusion. Many gabbroic and leucocratic regions; i.e., Sangun-Renge Belt (around 300 blocks with a few meters in diameter occur in the Ma), Suo Terrane (around 220 Ma) and Chizu Ohsa-yama body. These blocks consist of fine- Terrane (around 180 Ma). Ophiolitic rocks such grained gabbro, jadeitite, albitite and rodingite, as serpentinite and metagabbro occurring in the and have been found only within the ultramafic glaucophanitic terrane are also different in age body. Therefore, they are thought to have been from one region to another, though they are included into or formed within the ultramafic older than the surrounding crystalline schists in body before its emplacement into the present po- each region. For example, the Ohsa-yama ultra- sition in the surrounding schists. mafic body is included in the 180 Ma Chizu Ter- The Ohsa-yama ultramafic body is in fault rane, and a block of basic schist in this body has contact with crystalline schists on the north and a K-Ar age of ca. 250 Ma (Watanabe et al., 1987). east of the body. The crystalline schists are com- Also, the Tari-Misaka, Ashidachi and Ochiai- posed dominantly of pelitic and basic rocks with Hokubo ultramafic bodies belong to the Chizu small amounts of intercalated thin beds of Terrane, so that we infer that there exists a close psammitic and siliceous rocks. Their schistosity petrogenetic link among these ultramafic bodies. planes trend N-S or NE-SW and dip 30 to 60 ° NW. On a macroscopic scale, however, these I11. General Geology of the Ohsa-yama Area schists form a large fold with its axial plane The Ohsa-yama ultramafic body is exposed trending E-W and extending from the in an area around the summit of Mt. Ohsa, which Katsuyama through Ohsa-yama to Hokubo areas Petrography of primary peridotites from the Ohsa-yama area 133031'30E W Alluvium Granite wm Rhyolite Conglomerate Pelitic schist Wl Basic schist Ultramafic rock 3504'N Fig. 2 Geological sketch map of the Ohsa-yama (Mt. Ohsa) area (Mitsuno and Sugita, 1980). The Ohsa-yama ul- sion (Nozaka, 1987). The minerals formed by the tramafic body is situated at the crest of this fold. contact metamorphism include orthopyroxene, The unusual shape of the Ohsa-yama body may olivine, talc and tremolite, and the recrystallized result from bending associated with the large- Ohsa-yama rocks show distinct, granoblastic or scale folding. poikiloblastic textures. The ultramafic rocks oc- Metamorphic minerals composing the curring in the eastern portion of the Ohsa-yama crystalline schists adjacent to the Ohsa-yama body are not affected by the contact metamor. body include muscovite, chlorite and epidote in phism. Although these rocks have suffered in- pelitic schists, and chlorite, epidote, actinolite, tense serpentinization, they still locally preserve albite, stilpnomelane, muscovite and locally the original textures and contain olivine, glaucophane in basic schists. A lawsonite-bear- orthopyroxene, clinopyroxene and brown spinel ing basic schist has also been reported from the as relict minerals. Any layered structures on Ohsa-yama area (Hashimoto and Igi, 1970). Oc- both mesoscopic and microscopic scales are not currence of these metamorphic minerals sug- recognized in the Ohsa-yama body gests high P/T metamorphism in this area. Olivine, the most abundant relict phase, is The Ohsa-yama body and the surrounding equidimensional in shape and a few millimeters schists are intruded by a biotite granite and in diameter. Pyroxenes have commonly exsolu- hornblende-biotite granite, and overlain by tion lamellae; i.e., clinopyroxene lamellae in or- unmetamorphosed conglomerates and rhyolites. thopyroxene, and vice versa. Brown spinel has These younger igneous and sedimentary rocks an equidimensional shape in a serpentine ma- are Cretaceous in age (Mitsuno and Sugita, trix, and a vermicular shape within or in contact 1980). with clinopyroxene grains. The relict minerals are never in direct con- IV. Petrography tact with each other due to replacement by min- The Ohsa-yama ultramafic rocks have suf- erals produced during serpentinization. Gener- fered contact metamorphism by the granitic in- ally, olivine and orthopyroxene are replaced at trusion. The metamorphic aureole developed their rims and along cleavages by serpentine and from the southwestern margin to middle portion fine-grained magnetite, and spinel is fringed of the Ohsa-yama body along this granitic intru. with such opaque minerals as ferritchromite and Table 1 Representative micropobe analyses of relic minerals in the Ohsa-yama ultramafic body. Rock type Harzburgite Dunite Phase Olivine Orthopyroxene** Clinopyroxene Spinel Olivine Spinel SiO2 40.80 40.79 55.35 55.42 52.35 53.37 0.03 0.01 40.56 40.53 0.04 0.03 TiO2 0.04 0.01 0.00 0.01 0.02 0.00 0.05 0.05 · Al203 2.72 2.83 2.78 1.84 27.45 28.46 28.85 28.55 ， Cr203 0.80 0.81 1.09 0.82 40.03 40.04 38.64 38.26 FeO* 8.99 9.03 5.68 5.76 1.94 1.53 21.00 17.80 8.39 8.72 21.94 22.94 MnO 0.10 0.13 0.14 0.16 0.08 0.07 0.42 0.35 0.11 0.11 0.55 0.68 NiO 0.36 0.38 0.08 0.11 0.04 0.04 0.11 0.06 0.44 0.42 0.06 0.07 MgO 50.24 49.98 33.35 33.80 16.43 17.14 10.94 13.01 50.93 50.67 9.58 8.56 CaO 0.00 0.01 1.63 1.07 24.71 25.07 0.02 0.03 0.00 0.03 0.00 0.01 Na20 0.00 0.01 0.06 0.04 Total 100.49 100.32 99.79 99.98 99.48 99.93 100.02 99.76 100.43 100.48 99.71 99.15 5 0 4 4 6 6 6 6 32 32 4 4 32 32 NOZAKA Si 0.992 0.994 1.921 1.917 1.919 1.943 0.007 0.002 0.986 0.986 0.010 0.007 0.001 0.000 0.000 0.000 0.004 0.000 0.009 0.009 Al 0.111 0.115 0.120 0.079 7.937 8.102 8.362 8.375 and 0.022 0.022 0.032 0.024 7.765 7.647 7.513 7.529 Cr Tsu Fe 0.183 0.184 0.165 0.167 0.059 0.047 4.309 3.596 0.171 0.177 4.512 4.775 gio Mn 0.002 0.003 0.004 0.005 0.002 0.002 0.087 0.072 0.002 0.002 0.115 0.143 0.001 0.012 0.009 0.008 0.012 0.014 SHI Ni 0.007 0.007 0.002 0.003 0.001 0.022 Mg 1.822 1.816 1.725 1.743 0.898 0.930 4.001 4.685 1.846 1.838 3.512 3.176 Ca 0.000 0.061 0.040 0.971 0.978 0.005 0.008 0.000 0.001 0.000 0.003 ATA 0.000 Na 0.000 0.001 0.004 0.003 Total 3.006 3.004 4.012 4.013 4.006 4.007 24.137 24.124 3.014 3.012 24.045 24.031 Mg/Mg+Fe 0.909 0.908 0.913 0.913 0.938 0.952 0.481 0.566 0.915 0.912 0.438 0.399 Wo 3.1 2.1 50.4 50.0 En 88.4 89.4 46.6 47.6 Fs 8.5 8.6 3.1 2.4 A*** 0.494 0.504 0.523 0.524 Cr*** 0.483 0.476 0.470 0.471 Fe3+*** 0.023 0.021 0.007 0.005 *total ironasFeO. **Beam diameter: 30 microns. *** Al-Cr-Fe3+ atomic ratios in spinel. Fe3+ is caluculated from structural formula. Petrography of primary peridotites from the Ohsa-yama area magnetite. In some cases where serpentiniza- were selected for chemical analysis of constituent tion has proceeded completely, serpentine and minerals; one is dunite containing olivine and associated minerals form pseudomorphs after brown spinel (sample no. 03141083) and the olivine and orthopyroxene, exhibiting mesh and other is harzburgite containing olivine, orthopy- bastite textures. In contrast with olivine and or- roxene, clinopyroxene and brown spinel (sample thopyroxene, clinopyroxene tends to better pre- no. 02231183). Their representative analyses serve its original crystal shape. It suggests that are listed in Table 1. clinopyroxene is more resistant to serpentiniza- 1. Olivine tion than olivine and orthopyroxene. Olivine crystals in both the dunite and In spite of its stronger resistance to harzburgite are homogeneous in composition serpentinization, clinopyroxene occurs less fre- within individual grains and from grain to grain; quently than olivine and orthopyroxene. Spinel e.g., the Fo contents are restricted from 90.8 to is also a relatively resistant phase and in fact it 91.0 in the harzburgite sample, and from 91.2 to occurs more frequently than the other relict 91.5 in the dunite with the exception of a small phases, but is only a small amount in all speci- grain (Fo 92.6) that is included in a larger ferrit- mens examined. Therefore, clinopyroxene and chromite grain. NiO contents in olivines from spinel appear to have been a minor constituent of the Ohsa-yama body are plotted against Fo con- the primary peridotites. Abundance of the relict tents in Fig. 3 along with ranges of olivine com- minerals and their pseudomorphs indicates that positions from peridotites in the neighboring ul- the dominant primary rock types of the Ohsa- tramafic bodies reported by Arai (1980) and yama body have been dunite and harzburgite. Hamada (1982). All the plotted points of the Ohsa-yama olivine compositions are entirely in- V. Mineral Chemistry cluded in the compositional area for the Tari- All analyses in this study were carried out Misaka and Ashidachi olivines. by using a JEOL electron probe microanalyzer 2. Orthopyroxene (Model JXA-733) at Okayama University. The Orthopyroxene grains examined have in- correction procedure was that of Bence and Albee variably many thin exsolution lamellae of (1968) with alpha factors of Nakamura and clinopyroxene. The orthopyroxene analyses ob- Kushiro (1970). tained by a focused, narrow electron beam of 1 Two samples from the Ohsa-yama body μm or so in diameter show considerably disper- 0.6 (wt% Ohsa-yama ●Harzburgite + Dunite Tari-Misaka Z Ashidachi Ochiai-Hokubo 排 0.2 85 06 95 Fo (mol%) Fig.3 Correlation between Fo and NiO contents of olivine in harzburgite and dunite from the Ohsa-yama ultramafic body. Compositional areas for olivines in peridotites from the Tari-Misaka and Ashidachi bodies (data from Arai, 1980), and from the Ochiai- Hokubo body (data from Hamada, 1982) are also shown for comparison. 6 Toshio NoZAKA and Tsugio SHIBATA 3 2 En 5 10 Fs (mol%) Fig. 4 Orthopyroxene composition in the pyroxene quadrilateral and correlation between Cr 203 and Al2O3 contents. Symbols are the same as 0 0.5 1 1.5 those in Fig. 3. Cr2O3 (wt%) B 2 En Fig. 5 Clinopyroxene composition in the pyroxene quadrilateral and correlation between Cr 203 0 0.5 1 1.5 and Al2O3 contents. Symbols are the same as those in Fig. 3. Cr203 (wt%) sive values of Ca content within each grain, even 3. Clinopyroxene Analyses of clinopyroxene were carried out tered electron images. It is inferred from this by a focused, narrow electron beam, because the fact that the effect of clinopyroxene lamellae ad- exsolution lamellae in this phase are not so abun- jacent to or beneath the analyzed points was not dant as in orthopyroxene. Figure 5 shows that avoided completely.I Hence all quantitative the clinopyroxenes from the Ohsa-yama body analyses of orthopyroxene were carried out by means of a broad beam of 30 μm in diameter. The Al and Cr contents are almost the same as those results so obtained for the harzburgite sample of the Tari-Misaka and Ashidachi bodies, but are shown in Table 1 and Fig. 4. Although the their Wo contents are slightly higher than those Ohsa-yama orthopyroxenes appear to be rich in of the Tari-Misaka and Ashidachi bodies as well Wo component owing partly to the analytical pro- as the Ochiai-Hokubo body. This compositional cedure, they are rather similar in composition to difference may be accounted for by difference in those in the Tari-Misaka and Ashidachi equilibration temperature among the ultramafic harzburgites than the Ochiai-Hokubo peridot- bodies, but such a possibility cannot be examined ites. further because precise composition of orthopy- Petrography of primary peridotites from the Ohsa-yama area A Cr/ 0.4 Fe3+ 0.2 Fig. 6 Correlation between Mg/Mg+Fe 2+ and Cr/ Cr+Al atomic ratios and Cr Al-Cr-Fe3+ atomic ratios of spinel. 0.5 0.7 Symboles are the same as those in Mg/Mg+Fe2+ Fig. 3. roxene coexisting with clinopyroxene in the analyses of the relict minerals derived from pri- Ohsa-yama body is not known, as mentioned mary peridotites have revealed the following; 1) above. mesoscopic and microscopic layerings do not or 4. Spinel rarely exist, 2) the primary peridotites consist The analyzed grains of brown spinel ex- mainly of olivine, or of olivine and orthopyroxene hibit an equidimensional shape in the dunite and a vermicular shape in the harzburgite. The words, the dominant rock types are dunite and spinel grains in the dunite have been replaced or harzburgite, and 3) the constituent minerals are mantled by opaque minerals thicker than those rather similar in composition to those ofthe Tari- in the harzburgite. Both the spinels possess al- Misaka and Ashidachi bodies than the Ochiai- most identical Cr/Cr+Al ratios as shown in Fig. 6. Hokubo body. From these facts it is concluded The spinel compositions are plotted nearly half- that the Ohsa-yama ultramafic body belongs to way between the Al and Cr apices in the Al-Cr- the massive group of the Arai's (1980) classifica- Fe3+ triangular diagram, and fall within the com- tion. position ranges of spinel from the Tari-Misaka and Ashidachi bodies. Also, the Ohsa-yama References spinels vary significantly in Mg/Mg+Fe2+ atomic Arai, S. (1980), Dunite-harzburgite-chromitite ratio, depending on the type of rock in which they complexes as refractory residue in the Sangun- occur. They are more magnesian in the Yamaguchi zone, western Japan. Jour. Petrol., 21, harzburgite than in the dunite (Fig. 6). This may 141-165. be simply due to the difference in primary spinel Bence, AE. and Albee, A.L. (1968), Empirical correc- composition between the harzburgite and dunite, tion factors for the electron microanalysis of sili- or otherwise it may reflect the difference in the cates and oxides. Jour. Geol., 76, 382-403. degree of replacement of spinels by the secondary Hamada, T. (1982), Geological and petrological studies opaque minerals (Nozaka, 1987). on the Onoro-yama ultramafic complex, Hokubo- cho, Okayama Prefecture. Unpublished thesis, Okayama University (in Japanese with English VI. Conclusions abstract). Observations of the ultramafic rocks in the Hashimoto, M. (1968), Glaucophanitic metamorphism field and under the microscope, and chemical 8 Toshio NOZAKA and Tsugio SHIBATA of the Katsuyama district, Okayama Prefecture, Japan. Jour. Fac. Sci., Uniu. Tokyo, Sec. II, 17, 99- 162. Hashimoto, M. and Igi, S. (1970), Finding of lawsonite- glaucophane schists from the Sangun metamorphic terranes of the eastern Chugoku province. Jour. Geol. Soc. Japan, 76, 159-160 (in Japanese with English abstract). Hiroshima Prefecture (1964), Geological map of Hiroshima Prefecture, Japan. 1: 200,000. Igi, S. and Abe, K. (1969), Ultrabasic rocks in the east- ern part of the Chugoku Zone, Japan. Bull. Geol. Surv. Japan, 20, 39-50. Igi, S. and Sakamoto, T. (1977), Quadrangle series 1: 50,000, Geol. Surv. Japan. Mitsuno, C. and Sugita, M (1980) Geological map of Okayama Prefecture, Japan. 1: 100,000. Naigai- chizu. Nakamura, Y. and Kushiro, 1. (1970), Compositional relations of coexisting orthopyroxene, pigeonite and augite in a tholeiitic andesite from Hakone volcano. Contrib. Mineral. Petrol., 26, 265-275. Nishimura, Y. (1990), "Sangun metamorphic rocks": terrane problem. In Pre-Cretaceous terranes of Japan (ed. K. Ichikawa et al.), Pub. IGCP Project 224, 63-79. Nishimura, Y. and Shibata, K. (1987), Tectonic frame- work of the "Sangun metamorphic belt". High Pressure Belt of Inner Zone, 4, 45-52 (in Japanese). Nozaka, T. (1987), Contact metamorphism of serpentinites in the Ohsayama area, Okayama Pre- fecture. Unpublished thesis, Okayama University. Research Group of Peridotite Intrusion (1967), Ultrabasic rocks in Japan. Jour. Geol. Soc. Japan, 73, 543-553. Watanabe, T., Nishido, H. and Nagao, K. (1987), Additional data of K-Ar age of Sangun meta- morphic rocks and related tectonics. High Pres- sure Belt of Inner Zone, 4,11-12 (in Japanese).
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Nozaka and Shibata (1994) - Petrography of primary peridotites from the Ohsa-yama area, Okayama Prefecture.txt
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Gravity-Based Fault Mapping:
The Ishikari Lowland, Hokkaido, Japan
A. Yamamoto
Institute of Seismology and Volcanology, Graduate School of Science, Hokkaido University,
N10-W8, Kita-Ku, Sapporo, 060-0810, Japan
Abstract. This paper examines gravity structures
of the Ishikari Lowland of Hokkaido, Japan, by fo-
cusing relief-shaded Bouguer gravity, specifically to
relocate the faults (the Ishikari Teichi Touen FaultZone, ITTFZ), and to present gravity-based fault
mapping as another approach for estimate of fault
segmentation. Gravity analyses are strictly basedon dense gravity data measured by various institutes
and organizations. Bouguer relief is produced by il-luminating the light from eight directions to effec-
tively display the detailed gravity features varying
laterally along the azimuthal direction. A striking
linearity of the relief-shaded Bouguer gravity along
the ITTFZ is found on most of the relief maps, par-
ticularly on the maps for the azimuth of the due
east and west directions. In the central part of the
ITTFZ, however, the lineament in the relief-shaded
Bouguer map does not bear a good correlation with
the known fault distributions. In addition, the grav-
ity relief for the azimuth of the south-west direction
exhibits a remarkable lineament, extending south-
ward from near Bibai, whose southward continua-
tion can be traced to the south-east until 20~25 km
south of Atsuma along the westernmost boundaries
of pre-Neogene volcanics. This implies that the ge-
ometry of the southern end of the ITTFZ provides
distinct continuity along the relief-shaded Bouguer
lineament roughly to the south-east.
Keywords. Gravity anomaly, Fault mapping,
Ishikari Plain, Ishikari Lowland, Hokkaido, Sapporo,
Japan
1 Introduction
The Ishikari Plain is chief industrial region and
characterized by the largest alluvial lowland (the
Ishikari Lowland) in Hokkaido, Japan. Fig-1
shows a 3-D perspective view of topography withknown faults (Fig.la), and a simplified geologicmap (Fig.lb) around the Ishikari Lowland, where
Bouguer anomaly isolines are superimposed. Ge-
ology information in Fig.lb is based on the digi-
tal version of
1:1,000,000 scale geological map of
Japan (3rd edition) by Geological Survey of Japan
(1995). Thick colored lines and dotted red lines in
Fig.lb denote known active faults by Nakata and
Imaizumi (2002) and the Research Group for Active
Faults of Japan (1991), respectively. Thick color-
coded faults in Fig.l indicate, red: fault certainlyexists and its location is accurately determined, ma-
genta: fault certainly exists but its location is not
accurate, green: fault possibly exists but invisible,
and blue: estimated fault lying at depths, respec-
tively. As shown in Fig.l, the Ishikari Plain is char-
acterized by the largest late-Cenozoic lowland (the
Ishikari Lowland) in west Hokkaido. Wide-spread
deposition of alluvial deposits constitutes a large
part of the overall sedimentary sequence around the
plain. Fault-bounded plains or basins often accom-
pany a variety of tectonic motions such as reversefaulting. Specifically in Hokkaido, these motions
are associated with the plate movements and the
secondary tectonics. In the eastern margin of this
lowland lies the Ishikari Teichi Touen Fault Zone
(ITTFZ) which borders on the Miocene hill belts(Iwamizawa, Kurisawa and Umaoi hills). The IT-
TFZ has en echelon features and runs nearly in theN-S direction, bending toward west, and is classifiedas reverse fault system. (Research Group for Ac-
tive Faults of Japan, 1991; Oka, 1986; Oka et al.,
2001). The gravity field prevailing over the Ishikari
Plain is best reflected on the Bouguer anomaly map.
Recently, Yamamoto (2003a,b) gave preliminary re-
sults of Bouguer anomalies over the plain and dis-
cussed about their important features. This paper
examines gravity structures of the Ishikari Lowland
by focusing Bouguer gravity reliefs, specifically (1)to relocate the faults (the ITTFZ), comparing the
242
new active fault distributions (Nakata and Imaizumi,
2002) with the older ones (Research Group for Ac-
tive Faults of Japan, 1991), and (2) to present the
gravity-based fault mapping as another approach for
correct estimate of fault segmentation.
2 Gravity Data
In total, about 25,000 gravity data from Hokkaido
University, Geological Survey of Hokkaido (here-
after referred to as GSH), the Japan Petroleum
Exploration, Co. Ltd. (hereafter referred to as
JAPEX), etc. were collected for gravity study
around the plain. Particularly, gravity data mea-
sured by JAPEX number more than 20,000 in the
lowland and its vicinity with an average spacing of
300~400 m. Komazawa et al. (1998) published a
Bouguer anomaly map around the Ishikari Plain as-
suming a reduction density of 2.3 gjcrr? on the ba-
sis of land gravity data by JAPEX. In the present
study, gravity values for land stations obtained by
JAPEX were first recalculated to determine absolute
values based on the Japan Gravity Standardization
Net 1975 (JGSN75, Suzuki, 1976). Then, JGSN75-
based values were converted to the gravity valuesbased on the Japan Gravity Standardization Net
1996 (JGSN96, Nakai et al., 1997; Yamaguchi et al.,
1997). Land gravity data by GSH were compiled and
also recalculated, referring to the values of JGSN96,
so as to fit the reference gravity data by HokkaidoUniversity to which the absolute gravity values were
already assigned. Gravity data by Hokkaido Univer-
sity were obtained by a LaCoste & Romberg Model
G land gravity meter and a Scintrex gravity meter.
These data were fully corrected for latitudinal and
elevation effects, instrumental height, earth tide, and
secular drift of the spring. These gravity data were
then reduced for free-air and Bouguer corrections,
as well as for terrain effects, where spherical terraincorrections were applied to a radius of 80 km ac-cording to the method by Yamamoto (2002). In all
reductions the earth's sphericity is taken into consid-
eration. After an overall revision and reduction for
these gravity data, a Bouguer anomaly map, a resid-
ual anomaly map, and a first horizontal derivative
map of the region were constructed.
3 Relief-Shaded Bouguer Gravity
As shown in Fig.lb, gravity field over the IshikariPlain is characterized by distinct low anomaly as
compared to its surroundings. This gravity low is
mostly associated with significant Quaternary sed-iments spreading extensively over the area. Steep
gravity gradients are observed along the ITTFZ in
the eastern margin of the plain. They are trend-
ing almost in the N-S direction, which is closely cor-related with that of the active fault distributions.
While gravity anomalies around the Nopporo Hill
(NH), where the Nopporo Hill Fault Zone (NHFZ
in Fig. la) is newly identified as an active fault sys-
tem in the newly-compiled fault mapping (Nakata
and Imaizumi, 2002), are small in amplitude andare associated with no abrupt gravity changes com-
pared to its topography and surrounding fault dis-tributions, implying that no sharp density contrast
are appreciable at depths. To clearly show the grav-
ity structure around the Ishikari Lowland, particu-larly the sharp and fine linearity corresponding wellto the ITTFZ, we discuss here relief-shaded gravity
anomaly illuminated by the light from various direc-
tions. In general, Bouguer anomaly can be shaded
by giving illumination to highlight gravity structureslaterally varying along the azimuthal direction (the
direction of illumination). Therefore, gravity struc-
tures with lateral variations along the azimuthal di-rection display striking contrast and also show strik-
ing difference as compared with those with lateral
variations perpendicular to the azimuthal direction.
This means that gravity structures with lateral vari-
ations perfectly perpendicular to the azimuthal di-
rection never appear in gravity relief map shown in
shades of white and black. We first take a horizon-tal derivative of non-filtered Bouguer anomaly along
the azimuthal direction. As is well known, horizon-tal gradients of gravity anomalies have a multiresolu-
tion interpretation based on Poisson kernel (Green's
function for potential fields). Then we normalizethe derivative such that their values range from -1
(black) to 1 (white). This procedure makes brighter
the areas whose Bouguer values increase along the
illumination vector, while the areas with decreasing
values along the vector are shown gloomily (black-
ish color in the figure). Note that we simply change
the azimuth of view point and use the fixed eleva-
tion of zero degree (horizontal view) in this study.
In this way relief-shaded Bouguer gravity maps wereproduced for eight directions of illumination. Fig.2a
illustrates the relief-shaded Bouguer gravity map in
a planimetric view illuminated by the light from the
south-west direction. This is quite effective to makeclear the laterally heterogeneous gravity structure
varying along SW-NE direction around the ITTFZ.
Note that the illumination vector is shown on the
upper-right side of the figure. Fig.2b is the same as
Fig.2a, but known active faults are superimposed in
243
bright colors. Heavy colored lines and dotted red
lines in Fig. 2b depict known active faults by Nakata
and Imaizumi (2002) and the Research Group for
Active Faults of Japan(1991), respectively. Col-
ored faults in Fig.2b indicate, red: fault certainly
exists and location is accurately determined, ma-genta: fault certainly exists but location is not ac-
curate, green: fault possibly exists but invisible, and
blue: estimated fault lying at depths, respectively.
Characteristic features in Fig. 2 are summarized be-
low.
(a) A pronounced linearity of shaded relief along
the ITTFZ is uncovered. This lineament extends
from at least 5 km north of IWZ (Iwamizawa)
to several kilometers west portion of HKT
(Hayakita) and ATM (Atsuma), and also shows
a good agreement with distributions of the old
(dotted red line) and new (thick colored line)
active faults in northern and southern parts of
the ITTFZ (Fig.lb). Whereas, in central part
near NNM (Naganuma), the lineament of relief-
shaded gravity does not agree with the trend of
active faults.
(b) Two high gravity belts, running almost in the N-
S direction from near Iwamizawa (IWZ) to nearHayakita (HKT), can be found distinctly. These
striking two belts are depicted as a pair of white
and black shades which is protruding westward.Note that each of these two belts consists of a
pair of Bouguer high and low extremes. Both of
these two belts develop and are mapped in the
western margin of the Miocene hill belts.
(c) In contrast, gravity low is clearly observed along
the new active fault extending northward from
near KYM (Kuriyama) to near BBI (Bibai). This
feature can be stressed particularly on the gravity
relief map for the azimuthal direction of 0° (due
north) or 180° (due south) from north clockwise.
This low anomaly runs along the eastern margin
of the Iwamizawa Hill (IH) and the Kuriyama
Hill (KH). However, it does not continue south-
ward from KYM, implying that a sharp gravity
change does not extend toward south along the
eastern margin of the Umaoi Hill (UHN, UHC,
UHS).
(d) The southernmost extension of the lineament de-
scribed in (a) reaches about 5 km south of ATM
(Atsuma) with good and unbroken continuity.
This fact suggests that active fault zone alongthe ITTFZ is spatially distributed over a dis-
tance of more than 80 km, whose southward ex-
tension can be mapped until about 20 km southof ATM along the boundaries of pre-Neogene vol-canic rocks.
The fact (d) is the most intriguing result in this
study since the southward extension of the ITTFZ
can be well defined and mapped by the present grav-
ity relief analysis, where no noteworthy faults are dis-
tributed. From this result we infer that the geometryof the southern end of the known fault system (theITTFZ) provides distinct continuity over a distance
of 20~25 km along the relief-shaded Bouguer linea-
ment roughly to the south-east. Consequently, ourresults suggest that the relief-shaded Bouguer grav-
ity analysis, using closely spaced Bouguer data, is
quite effective in demarcating the extension of thepre-existed (or invisible) faults for correct estimates
of their location and segmentation.
4 Conclusions
Bouguer gravity relief was produced by illuminat-
ing the light from eight directions to effectively de-pict the detailed gravity features varying laterally
along the azimuthal direction. A conspicuous linear-
ity of shaded relief along the ITTFZ is detected from
Bouguer anomaly relief particularly on the map for
the azimuth of the south-west direction. This lin-eament extends from at least 5 km north of IWZ
(Iwamizawa), ending over several kilometers west of
HKT (Hayakita) and ATM (Atsuma). Also, this lin-eament shows a good agreement with distributions of
the old and new active faults in northern and south-
ern parts of the ITTFZ. Furthermore, a remarkable
lineament extending southward from near BBI to
south of ATM is found on gravity relief map partic-
ularly for the azimuthal direction of 0° (due north)
or 180° (due south) from north clockwise. This factasserts that the active fault strands along the ITTFZ
is spatially distributed over a distance of more than
80 km and its southward extension can be mappeduntil 20~25 km south of ATM, separating the pre-
Neogene volcanic rocks. Concludingly, we note that
relief-shaded Bouguer gravity analysis plays an im-portant major role in the gravity-based fault map-
ping to locate and characterize the various strands
and segmentation of a fault zone.
Acknowledgments. The author gratefully acknowl-
edges the personnel at the Japan Petroleum Exploration,
Co. Ltd., who kindly permitted to use their land gravity
data. This paper benefited from constructive criticism
made by J. A. R. Blais. All figures were produced using
the GMT (Generic Mapping Tools) software of Wessel
and Smith (1998).
244
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ppl57, (in Japanese).
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tive faults in Japan, sheet maps and inventories, re-vised edition, University of Tokyo Press, pp437, (in
Japanese).
Suzuki, H. (1976). The International Gravity Standard-
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ment of the Japan Gravity Standardization Net 1996,
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1996, convened and edited by J. Segawa, H. Fujimoto
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gravity anomaly using a digital elevation model grid-ded with nodes at every 50 m, J. Fac. Sci., Hokkaido
Univ., 11, No.6, pp.845-880.
Yamamoto, A. (2003a). Gravity anomaly atlas of the
Ishikari Plain and its vicinity, Hokkaido, Japan,
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Yamamoto, A. (2003b). Gravity-based active fault map-
ping around the eastern margin of the Ishikari Low-
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12, No.l, pp.17-39, 2003.
A. Yamamoto, Institute of Seismology and Volcanol-
ogy, Hokkaido University, Sapporo, 060-0810, Japan.
(Email:star@eos. hokudai.ac.jp)
245
fa)
7/
I Jf •?•-. • Jr-m t
Fig. 1. (a) A 3-D perspective map showing topography and known faults around the Ishikari
Lowland. Colored active faults are taken from Nakata and Imaizumi (2002). Vertical ex-
aggeration is about 5 times. ITTFZ: the Ishikari Teichi Touen Fault Zone, IH: Iwamizawa
Hill, KH: Kurisawa Hill, UHN, UHC, UHS: Umaoi Hill (Northern, Central and Southern
Block), NHFZ: the Nopporo Hill Fault Zone. (b) Simplified geologic map around the Ishikari
Lowland. Bouguer anomaly isolines are superimposed with a contour interval of 5 mgal. As-sumed density is 2.67
g/cm3. Geology information is taken from Geological Survey of Japan
(1995). Heavy colored lines and dotted red lines demonstrate known active faults by Nakata
and Imaizumi (2002) and the Research Group for Active Faults of Japan (1991), respectively.
Large closed triangles and squares, followed by three letters, show geographical locations ofmajor named summits and cities (towns), respectively. NH: Nopporo Hill, SPR: Sapporo.
246
(a)
(b)
(3
Fig. 2. (a) Relief-shaded Bouguer gravity map of the Ishikari Lowland and surrounding
areas, illuminated by the light from the south-west direction (azimuth: N135°W, elevation:
0°). Azimuthal direction is shown on the upper-right side of the figure. No filtering is ap-
plied to Bouguer anomaly. (b) Same as (a), but known active faults are superimposed in
bright colors. Heavy colored lines and dotted red lines denote known active faults by Nakata
and Imaizumi (2002) and the Research Group for Active Faults of Japan(1991), respectively.Colored faults indicate, red: fault certainly exists and location is accurately determined,
magenta: fault certainly exists but location is not accurate, green: fault possibly exists butinvisible, and blue: estimated fault lying at depths, respectively.
247
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Yamamoto (2005) - Gravity-based fault mapping the Ishikari lowland.txt
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Geothermics 39 (2010) 228-241 Contents lists available at ScienceDirect Geothermics ELSEVIER journalhomepage:www.elsevier.com/locate/geothermics Temperature and chemical changes in the fluids of the Obama geothermal field (SW Japan) in response to field utilization Hakim Saibi*, Sachio Ehara LaboratoryofGeothermics,DepartmentofEarthResources Engineering,FacultyofEngineering,KyushuUniversity,744Motooka,Nishi-ku,Fukuoka19-0395,Japan ARTICLEINFO ABSTRACT Article history: Thermal waters from Quaternary volcanic rocks (predominantly andesites) discharge along faults in the Received 20 February 2009 Obama geothermal field of southwestern Japan. The chemistry of more than 100 thermal and ground Accepted 8 June 2010 water samples collected between 1936 and 2005 indicate that the Na-Cl hot spring waters are a mixture of "andesitic" magmatic, sea and meteoric waters. Mixing models and silica and cation geothermome- try were used to estimate the SiO2 and Cl composition and the temperature (~200°C) of the reservoir Keywords: fluids deep in the geothermal system. The isotopic data (18O and D) are consistent with a mixed origin Geothermal interpretation of the waters feeding the Obama hot springs, i.e. a large proportion of meteoric and sea Hydrochemistry waters, and a small magmatic component. Temperatures and chemical concentrations of the thermal Geothermometry Obama geothermal field waters were affected by the 1944-1959 salt production operations, but have recovered after closure of Japan the salt factories; now they are similar to their pre-1940 values. In the future, the Obama geothermal field may be suitable for electric power generation, although heat and fluid extraction will require careful @ 2010 Elsevier Ltd. All rights reserved. 1. Introduction an average depth of about 100 m. With the help of pumps they extracted about 750-1800 m3/day of hot fluids from the system; The Obama geothermal field (OGF) is located on the Shimabara between 1800 and 2400 m3 of hot water was needed to produce Peninsula, west of the Unzen volcano, in the western part of Kyushu 1 t of salt. All the geothermal wells were located very close to the Island, southwestern Japan (Fig. 1). Spas fed by the Obama geother- 99 3.2 factories, artesian hot springs and sampled hot springs. high temperature (IHES, 2002); thermal waters discharging along As a result of the over-pumping that accompanied commercial s no e s r e i ars salt production, the Obama geothermal aquifer was increasingly The hot spring waters of the OGS have been classified as being of affected by seawater intrusion, which raised the salinity and the NaCl-type (Saibi et al., 2006a); the proximity of the geother- decreased the temperature of the waters to the point that boiling mal field to the coast facilitates the intrusion of seawater into the ceased in the wells. By 1950, the lack of boiling wells was becom- hydrologic system (Saibi et al., 2006b). 0.4 Watanabe (1958) noted that there were 76 wells in the OGF and running dry. despite its small surface extension (around 1.5 km?) the hot water In 1951 there were 84 hot springs, 14 of them artesian. In 1955 output was very large, i.e. 56,500 t/day. Later, Yuhara et al. (1986) the number of salt factories was reduced to 40. At that time, only measured a total flow rate of 8168 t/day from 22 wells in the area. 277 Production of salt (NaCl) through evaporation, enhanced by dissolved solids (TDS) in the water reached 25,000 mg/L, which was heat from the thermal waters, began in 1944 and has played an three to four times the concentration before salt production began. important role in the history of the OGF. The number of salt fac- 247 tories rapidly increased to 80 by 1951, at which time the total 32,000 m3/day in 1955, without impacting salt production levels. geothermal water extraction was about 50.000 m?/day. The annual The number of artesian hot springs more than doubled (from six salt output was 7600t in 1948, whilst it averaged ~10,000 t/yr in to 12) between 1955 (Fig. 2a) and 1988 (Fig. 2e), especially in the 6.4 northern part of the geothermal field. On 17 September 1959, Typhoon 14 struck the Obama Spa, caus- ing serious damage and leading to the abandonment of most of the salt factories by the end of that year. Also by that time, 53 hot E-mail address: saibi-hakim@mine.kyushu-u.ac.jp (H. Saibi). springs showed a decline in flow rate, with only 24 being appro- 0375-6505/$ - see front matter @ 2010 Elsevier Ltd. Allrights reserved. doi:10.1016/j.geothermics.2010.06.005 H.Saibi,S.Ehara/Geothermics 39(2010)228-241 229 a) urumiV Japan Kuiuy AsoV. EastChinaSea KirishinaV. KaimondakeV PacificOcean 50km 132 Obama geothermalfield ShimabaraPeninsula Hatchobaru geothermal field Ogiri geothermal field UnzenV Yamagawa geothermal field V.=Volcano (b) Fig. 1. Map showing the location ofthe Obama and other geothermal felds in Kyushu Island,Japan.(a) Black squares: geothermal areas related to Quaternary-Recent volcanic activity; black triangles: active volcanoes. (b) Black open rectangle: location of the Obama geothermal field and of the area shown in Fig. 4. priate for use in public baths, hotels and inns, which prompted hot affected by tectonic and volcanic processes (Tamanyu and Wood, water production to be reduced to 8150 m3/day. The last salt fac- 2003). The graben occurs at the junction of the Southwest Japan tory was closed in 1965. Because of diminished exploitation the Arc and the Ryukyu Arc, and was formed either by extension asso- geothermal resource gradually recovered, i.e. the number of boil- ciated with the subduction of the Philippine Sea Plate beneath e a sdn m s ei i n m the Eurasian Plate (Kamata and Kodama, 1993), or by right-lateral water salinities decreased. As a result, the present thermal spring shear caused by oblique subduction of the Philippine Sea Plate characteristics are similar to those observed before 1940. beneath Kyushu Island (Tsukuda, 1993). There are many Quater- The paper focuses on the main changes in hot spring nary volcanoes within the graben, including Mts. Tsurumi, Kuju, Aso water chemistry, and in estimated reservoir temperatures. and Unzen (from east to west), which are accompanied by active Chloride-enthalpy and silica-enthalpy diagrams, as well as cation geothermal systems (Fig. 1). A number of geothermal manifesta- and silica geothermometers were applied to hot spring water sam- tions are found along the faults bounding the graben. ples of the Obama geothermal system (OGS) collected and analyzed The Obama area is dominated by Quaternary volcanic forma- from 1936 to 1987, and in 2005, in order to investigate reservoir tions (Fig. 4); the volcanics in the area of the OGF are primarily andesitic. Based on petrographic and macroscopic analysis of drill resources ofthe OGF Isotopic data (18O and D) on waters from river, cuttings and core samples from well Uz-2, the subsurface stratig- wells and springs, helped to resolve the origin of the hot spring s uz s si o n waters, and to propose a new hydrological model for the Obama Kuchinotsu Group (Fig. 5), both of Quaternary age (NEDO, 1988). geothermal system. The Unzen Volcanics are found at the surface and to depths of about 500 m, and consist of hornblende andesite lavas and tuff 2. Geologic setting breccias of the Takadake and Obama Formations. The lavas are intercalated by tuffs and silts of the Tatsuishi Formation that occur The Obama geothermal area is located in the southwestern part in the area at two depth levels in well UZ-2, i.e. at 197-291 m, and of the Beppu-Shimabara graben (Fig. 1), a structural depression 447-512 m (Fig. 5). The reservoir in the OGF is mainly located in 230 H.Saibi,S.Ehara/Geothermics 39(2010)228-241 (a) 1955 400m (b) 1961 Chijiwa Bay Chijiwa Bay (C) (d) 1979 1987 Chijiwa Bay Chijiwa Bay (e) 1988 Chijiwa Bay ONo discharge (stopped) Discharge of hot water by using pump Artesian hot springs Location of salt factories Fig. 2. Location of hot springs and wells in the Oba Chijiwa Bay 1984 and 2001 500m Hot spring (1984) Hot spring (2001) well number (1984) Fig. 3. Location of hot springs and wells mentioned in Table 6. H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 231 130°12'25” 130°12'46" UZ-2 0m(depth) N Rain ple TakadakeFormation B Hornblende andesite lavas and tuff breccias wa Volcanics Fault 197.3m Tatsuishi Formation 20" Z-2(Well) silts and volcanic tuffs 3243' 291.2m Fault River water sample zen un ObamaFormation Hornblende andesite lavas 447m Tatsuishi Formation silts and volcanictuffs 511.6m Kita-arima Formation Tatsuishi Formation of Unzen Volcanics tuffs, sands and mudstones 536.5m (Volcanic breccias and tuff breccias) TakadakeFormation-Unzen Volcanics (Hornblende andesitic volcanic breccias) C Takadake Formation-Unzen Volcanics Minami kushi-yamaVolcanics pyroxene andesites, pyroclastics and basalts (Volcanic breccias) Lake deposits (Volcanic ashes, sands and silts) Takadake Formation-Unzen Volcanics 773m (Biotite hornblende andesites) 0 1 km Kita-arima Formation 1 tuffs, sands and mudstones Fig.4. Geologic map of the Obama geothermal field (see Fig. 1b modified from NEDO Gro (1988). The black squares indicate the positions of the sampled Obama hot springs. The location of this figure is shown in Fig. 1b (black rectangle). Kuchinotsu the Takadake and Obama Formations; the Tatsuishi Formation is of lower permeability. The Kuchinotsu Group, found between 512 1071.7m and 1500 m depth (Fig. 5), corresponds to the basement rocks. Oya Formation (upper part) 3. Water chemistry andesites, tuffs, sandstones, silts and mudstones 3.47 and rain waters collected from the Obama area. The composition of the OGF hot spring waters has been collated from several surveys 1255.7m including: OyaFormation(lowerpart) 1- Nagasaki Prefectural Institute of Health and Environmental Sci- andesites, tuffs, sandstones, silts andmudstones ences (IHES, 2002): 1936-1987 chemistry data; 2- New Energy Development Organization (NEDO, 1988): 1984 chemical and isotopic data; 1499.8m 2.98 Fig.5. Stratigraphic log for well UZ-2 (modified from NEDO, 1988); well location is 4- Kyushu Environmental Evaluation Association (KEEA, 2005): shown in Fig. 4. Wellhead elevation: 132 m above sea level. 2005 chemistry data. In the past, researchers estimated the chemical composition of the xa sse s s 0.96 excluded from the data set. ses (weight and volume), although the exact methods used were 4.81 0.66 in the waters of hot springs N.4 and N.26, which have been sampled mined the chemistry of the water samples. The ionic balance repeatedly; N.4 is located close to the coast, while N.26 is further provides an indication of the accuracy (and hence reliability) of the 11.6 analyses, which is given by, to varying degrees of mixing of the thermal waters with sea and ≥ Cations (mequiv./I) - ≥Anions (mequiv./) fresh ground waters, similar to what is observed in the Reykjanes, Balance(%) ×100 Iceland, geothermal field (Sveinbjornsdottir et al., 1986). The figure ≥ Cations(mequiv./1) + ≥ Anions (mequiv./1) shows that the waters of hot spring N.4 have higher temperatures 29963/10.9 0.7 126 30 151 21 84 10,630 2916 207 213 5371 564 136.94 18 1.113 127 156 202 141 131 34 155 19 4963/18.0 9g 12,630 3493 264 317 251 6404 635 154.64 22 8'0 1.111 130 163 208 150 140 40 164 18 8963/10.8 12,340 Kyushu, Japan. Res. Geol., 60, 359-376. 256 225 6300 603 156.17 21 0.4 1.110 130 164 207 151 140 41 165 18 22963/27.5 80 11,630 The Hishikari Deposit -, 14, 63-69. 250 308 231 6058 573 140.02 19 0.8 1.087 130 157 207 143 133 35 157 18 291963/27.5 83 10,260 2799 168 medium-K volcanic rocks in southern Kyushu, Japan, for the contribution of a felsic lower crust, Chem. 5239 523 145.40 17 0.1 1.114 120 160 193 146 135 37 160 21 4963/27.7 86 12,690 3600 263 334 255 6602 614 111.55 23 1.0 1.088 129 144 basalt magma as an explanation for the origin of 126 117 23 142 18 8963/27.7 6 12,560 3581 261 317 for contribution of deep crustal fluid to the Hishikari 240 6557 602 105.40 23 0.6 3380 1.088 129 140 20 206 123 114 20 138 18 22964/37.9 82 12,380 256 324 6270 600 136.94 0.7 1.118 130 156 207 141 19 0.0 131 34 155 18 291964/37.2 84.5 11,450 3090 214 301 224 5820 541 150.02 1.124 126 162 201 126 148 137 39 162 19 4964/37.9 98.5 12,500 3390 260 313 230 6350 601 110.78 22 0.1 1.122 131 143 208 117 22 141 18 8964/37.9 86 12,300 3380 252 228 6250 592 103.09 IZ 0.2 1.118 4.0 130 139 207 121 I12 18 137 18 26971/57.8 94 8740 2700 81 397.5 140.5 4491 748.5 220.5 5.69 1.003 18 49 18 101 10 -11 12 43 22975/12.7 9850 2780 200 5020 447 224 130.79 1.075 149 H 6 176 Z'0 153 215 128 31 152 14 Saibi 291975/129 94.1 8880 2520 258 195 158 4530 419 229 119.25 16 0.1 1.069 144 147 206 131 122 26 146 16 4975/12.9 98.5 Volcanol. Geotherm. Res., 8, 161-175 2720 273 163 5090 440 195 83.09 18 0.4 1.073 153 127 232 66 8 124 15 Ehara / Geothermics 8975/18.1 100 0886 2830 34 271 155 5060 417 159 79.24 18 0.1 1.107 65 125 124 104 97 6 121 35 26977/12.6 95 9158 2735 276 192.2 111.1 4723 567.4 239.2 124.94 8.6 0.5 1.036 147 150 205 134 125 29 149 15 22978/77.5 96.7 0056 2739 262 204 113 4913 192.7 206.2 228.49 14.3 1.0 1.101 144 190 Hishikari gold deposit, southwest Japan: implications 183 170 65 193 16 2979/18.1 99 9275 2788 194 192 4886 410 167.6 147.71 5.0 16.6 3.1 1.037 146 161 211 147 136 88 161 15 20979/12.8 98.1 9142 2679 278 192.5 195 4635 440 234.3 230.03 16.1 1.4 3.5 1.056 144 161 207 184 170 66 61 15 979/18.2 100.2 9596 2938 309.7 182.1 179.6 4931 412 74.3 184.02 6.0 17.2 4.9 1.063 146 175 208 164 152 51 176 14 18979/77.6 87.8 9182 2668 325 189.7 Assimilation of lower to middle crust by high alumina 4865 486.3 153.8 264.03 5.1 13.7 0.1 1.044 156 102 221 197 180 182 75 205 14 291979/77.5 79.1 8760 2553 281.3 185.3 105.3 4668 469.6 86 220.03 4.5 15.3 0.9 0.6 1.035 151 187 212 166 63 190 15 39(2010)2 4979/77.4 66 2622 296.9 61 105.3 4956 465 916 200.03 5.2 13.3 1.5 1.068 153 181 215 172 159 56 8 15 8979/77.6 92.2 9083 2630 337.5 181 104.4 4836 439 145.2 200.03 5.6 15.3 0.2 1.047 159 181 226 172 159 56 183 13 22983/67.9 94.5 8678 2540 260 215 130 4560 384 186 172.33 15.7 0.5 1.049 146 170 222 159 147 47 172 16 291983/67.4 88 2520 240 215 128 4400 374 260 200.03 16 1.3 1.041 143 181 218 172 159 8 17 228 8983/68.1 96.5 9060 2660 360 185 106 4780 364 186 140.02 17 0.1 1.048 161 157 231 143 133 35 157 13 -241 1986/18.3 84 9118 2714 600 165.5 136.6 4959 368.4 152.5 234.80 5.9 15.1 1.002 185 Z6 186 172 67 195 8 23987/37.6 16 9060 2543 267.9 170.1 126.3 4503 355.1 209.3 215.72 6'9 23.7 2.3 1.108 148 186 208 178 165 61 189 15 24987/37.9 93.8 8940 exchange in biotite. Am. J. Sci., 274, 396-413. 250 187.5 127.5 4608 364.9 183.7 197.18 6.4 22.6 1.8 1.078 144 180 202 170 158 55 182 16 28987/37.5 90 8510 2400 253 173.6 125.1 4326 345.4 238 205.26 6.1 22.4 1.6 1.083 147 182 208 174 161 58 185 16 29987/38.0 96 8792 Geochim. Cosmochim. Acta, 53, 269-278. 247 173.6 125.1 4361 350.2 192.2 202.80 6.5 21.6 2.2 1.112 145 182 204 173 160 57 8 16 10987/48.5 99.5 9270 Kyushu, Japan. J. Geochem. Explor., 36, 1-56. 310 173.2 203.8 4964 380 150.6 66.16 5.3 20.8 0.3 1.6 1.040 148 116 215 93 87 -2 111 14 12987/48.4 86 8930 2692 285.7 169.6 180.7 4822 376 137.8 69.24 5.0 22.2 0.4 1.7 1.031 146 118 209 96 89 0 114 14 131987/48.5 L6 8855 2653 309.5 166.1 184.6 4786 366 150.6 69.24 4.8 23.3 0.3 1.7 1.028 150 118 217 96 89 114 14 17987/48.3 96 9035 2615 273.8 173.2 165.4 4857 384 124.9 69.24 4.7 21.7 0.3 0.0 1.051 145 118 208 96 68 0 114 146.4 15 27987/47.6 69.5 3850 1107 108 91.1 107.6 2076 196 64.62 1.8 11 0.2 1.005 139 114 202 92 98 85 E- 110 20 4987/48.3 99.5 8845 2576 286 167.8 165.4 4716 141 72.32 4.6 22.3 0.2 1.050 148 120 213 91 2 116 128.1 214 14 7987/48.4 100.5 8815 2653 297.6 164.3 196.1 4786 360 73.86 5.4 23.8 0.3 2.0 1.027 147 121 100 6 117 14 8987/48.4 99.5 8735 2615 297.6 162.5 192.3 4716 350 131.3 69.24 5.0 24.5 0.3 1.9 1.032 148 118 215 96 89 0 114 91987/48.2 93.8 8765 2615 285.7 166.1 184.6 4751 356 112.1 72.32 5.1 23.3 0.3 1.8 1.035 146 120 211 86 91 2 116 14 1987/48.4 16 8790 2653 285.7 164.3 176.9 4786 068 133.3 69.24 4.7 21.8 0.4 1.5 1.023 146 118 210 96 68 114 14 2005/48.5 97.4 8354 2800 260 120 130 4500 180 5.1 16 4.4 1.004 142 198 14 %, ionic balance. Inferred reservoir temperatures.A: NaKCa geothermometer (Fournier and Truesdell, 1973); B: silica (Fournier and Potter, 1982): C: cation composition geothermometer (CCG)(Nieva and Nieva, 1987): D: silica 234 H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 Table 2 Chemical composition (in mg/L) of different types of waters in the Obama geothermal field (NEDO, 1988). The location where the river water was sampled is shown in Fig. 4. Type of water Date pH T(°C) Na K+ Mg2 Ca2+ Cl- SO42- HCO3 SiO2 Sea December 1975 17.8 10,250 358 1270 399 18,650 2450 138 1.3a River 1984 7.3 16.2 5.94 2.31 2.19 6.43 5.50 2.10 37.20 51.70 31/10/1999 7.0 17.8 5.80 2.20 2.50 5.70 5.40 5.10 36.40 40.10 a Sampled on 13/12/2000 (Ohsawa; pers. comm., November 2008). b Geological Survey of Japan (2010); unpublished data. (a) (b) TDS (mg/L) 98- 20,000 94- 06 15,000- 86- Temperature(°℃) 82- 10,000 口 口 78- 19281938 1948 1958 1968 1978 1988 19281938 1948 1958 1968 1978 1988 (d) 10,000- (C) .... Na o K →-·CI Concentrations Mg 。 Concentrations (mg/L) (mg/L) Ca 1000- SO4 7500- HCO< △ 5000- 500 2500- 口 100 192819381948 1958196819781988 1928193819481958196819781988 Year Year Fig. 6. Changes in the characteristics of the waters of hot spring N. 4 (black symbols and lines) and hot spring N. 26 (open symbols) with time. (a) Temperature changes; (b) changes in TDS; (c) changes in Na and Cl concentrations; (d) changes in K, Mg, Ca, SO4, HCO3 and SiO2 concentrations. and concentrations than N.26, except for SO4 and HCO3. In general, (1936-1940); Group 2 to the salt production stage (1948-1955); further inland the salinity of the thermal waters decreases, and is Group 3 to the period characterized by decreased salt produc- between 100 and 2500 mg/L. tion (1956-1959); and Group 4 to when the thermal springs were The pH of the higher temperature (52-102 °C) spring waters in mainly used for bathing (1960-2005). the OGF is in the 7-8.5 range. The alkalinity of these waters can be explained by the depletion of HCO3 (i.e. separation and loss of CO2) during boiling (Truesdell, 1995). The TDS of the hot (>40 °C) spring 20,000 waters range from 3850 to 25,900 mg/L. 18,000 Seawate The dominant ions in the hot spring waters are Cl and Na, Y=1.9X-226.2 with concentrations ranging from 2076 to 13,250 mg/L and 1107 16,000- R² = 0.99 to 7182 mg/L, respectively. Such values are commonly found in hot 14,000 springs fed by fluids that rise rapidly (i.e. directly) from deep reser- (7/6u voirs, and thus can be used to resolve the character of the upflows 12,000 E in geothermal systems (Cortecci et al., 2005). 10,000 At Obama, most of the chloride in the thermal water is likely 8000 to come from seawater, while a small amount may be derived 口Group1 from HCl degassing from a magma chamber located beneath the 6000- △Group 2 o s sn n ss p 4000- 0Group3] the sodium. The concentration of sulfate in the thermal waters is also high (7.8-1575 mg/L), and in part probably derives from the 2000 Group4 Riverwater oxidation of H2S (up to 12 mg/L in the NaCl springs; Ellis, 1969) or from seawater. 2000 4000 6000 8000 10,000 The chemical data on the Obama hot spring waters have Na (mg/L) been divided into four groups according to periods of contrasting Fig. 7. Sodium-chloride plot for thermal, sea and river waters from the Obama geothermal resource utilization: Group 1 corresponds to the ini- geothermal feld. Group 1: 1936-1944; Group 2: 1945-1955; Group 3: 1956-1959; tial stage when bathing was the primary use of the spring waters Group 4: 1960-2005. H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 235 1000- G 100 Riverwater Seawater Group1 1.0 Group2 Group3 Obama Seawater Group4 (Nozaki, 1992) 0.01 Mg Ca Na K CI SO4 BrHCO3 103x He 0.01310²xAr Fig. 8. Schoeller diagrams for the hot and cold waters of the Obama geothermal He/Ar field. Fig.9. Relative He, Ar and N2 contents in gas dissolved in Obama thermal well and sea waters (Nozaki, 1992) and in air (Ohsawa, 2006). Fig. 7 shows a plot of sodium vs. chloride concentrations in the Obama hot spring and representative sea and river waters. The fig- ure indicates that during the period of salt production (Groups 2 elemental ratios stabilized during the periods when Groups 3 and and 3) the Na and Cl contents in the thermal water increased, and 4 waters were collected. that since these operations ceased in 1959 they are returning to The high Br/I ratios observed even after the extraction of salt their original values (pre-salt production times; Group 1). was reduced and stopped could be explained by the contribution of We used chemical stoichiometry to calculate the concentrations organic matter in shallow near-surface reactions (Nicholson, 1993). of SiO2 in the water samples, based on the molecular weights of After the end of salt production the Br/Cl also remained high for H2SiO3 and SiO2. The SiO2 content in seawater collected in 2000 Groups 3 and 4 (Table 3) because the chloride contents in the hot from Chijiwa Bay near the Obama hot springs is 1.3 mg/L (Shinji waters ofthese two groups decreased drastically compared to those Ohsawa, pers. comm., November 2008). of Br. According to Fig. 8 the water samples collected in the Obama From the ternary Nz-He-Ar diagram (Fig. 9) it can be inferred area are ofthe alkali-chloride type. The figure also shows that chem- that the gases from the thermal water of OGF wells have a meteoric ical concentrations of the primary cations and anions present in the origin (Ohsawa, 2006). The Cl-SO4-HCO3 ternary diagram (Fig. 10) hot water discharges are generally lower than in seawater. proposed by Giggenbach (1988) indicates that all of the waters The Schoeller plots of Fig. 8 indicate that the Obama thermal d no p a ss n n waters are generally low in Br and HcO3, as are the cold river waters waters, but also are similar in composition to seawater. of the area. The thermal waters have high concentrations of Mg, Ca, The K-Na-Mg diagram (Fig. 11) (after Giggenbach, 1988) indi- Na, K, Cl and SO4, with seawater having higher amounts of Mg, Na, cates that none of the thermal spring waters are in full equilibrium Cl, and SO4. Analyses done in 1979-2005 show an average Lit con- with the host rocks and that they possibly have mixed with different tent of 5.19 mg/L in the hot spring waters (Table 1), reflecting the types of water (Barragan et al., 2001). predominantly andesitic composition of the host rocks (Nicholson, Fig. 12 shows the change of major cation and anion concen- 1993). trations in waters from selected Obama springs with time. They The Mg concentrations in the Obama thermal waters range from increased between 1936 and 1944, stabilized during 1944-1955, 19.8 mg/L (1937 samples) to 1045 mg/L (1956 samples) (Table 1) decreased from 1952 to 1970, and generally did not vary after- id un n s jo spuesnou on spapun ae p wards. Between 1936 and 1970 the changes in Cl and Na (0.01-0.1 mg/L; Nicholson, 1993). The high concentrations may indicate near-surface reactions, leaching of Mg from the host rocks 100 and also mixing with seawater, which has elevated Mg contents. Group1 The chemical characteristics of the Obama hot spring waters Group 2 indicate that most of them have a major seawater component 0Group 3 (Table 1). This is reflected, for example, by the high Ca (and to a Group 4 lesser degree, Mg) concentrations in the thermal waters, which is Seawater evidence of mixing with the seawater that intrudes into the shal- Riverwater lower parts of the geothermal field. Also, the B/Cl ratios in waters extracted from OGF wells in the 0.005-0.015 range (NEDO, 1988) Q confirm that mixing with seawater occurred (Giggenbach, 1992). Bromide concentrations are generally low in geothermal fluids, except when there is mixing with seawater (Nicholson, 1993). The analytical results (Table 3) show that the Br/Cl ratio in Obama spring waters is small for Group 1, but is greater for Group 2, which corre- sponds to period of increased resource exploitation. Subsequently, 100 Steam-heated/Steamcondensates Table 3 Br/Cl and Br/I ratios for each of the four spring groups. HCO3 100 Parameter Group 1 Group 2 Group 3 Group 4 Fig. 10. Relative Cl-SO4 and HCOs contents in thermal, sea, and river waters of Br/Cl 0.00132 0.00335 0.00325 0.00371 the Obama geothermal feld. Group 1: 1936-1944; Group 2: 1945-1955; Group 3: Br/1 0.41 23.47 81.24 51.44 1956-1959; Group 4: 1960-2005. 236 H. Saibi, S. Ehara /Geothermics 39 (2010)228-241 Na/1000 (a) Group1 Na 12,000 Group 2 CI 13 Full O Group 3 equilibrium 8000 Group 4 Seawater 4000 ： 1000 1930 1956 1982 2008 Immaturewaters K/100 /Mg 1600 (b) ·K Fig. 11. Ternary Na-K-Mg diagram mg/L for Obama hot spring waters (after 1400- . Mg Giggenbach, 1988). Group 1: 1936-1944; Group 2: 1945-1955; Group 3: Ca 1956-1959; Group 4: 1960-2005. 1200- SO4 1000 HCO 800 concentrations are greater than in SO4, Ca, and Mg, especially dur- 600 ing the 1945-1960 period. 400- The variations in spring water chemistry shown in Fig. 12 reflect 200- years. Higher levels of salt production or of thermal water extrac- +0 tion enhance seawater intrusion into the geothermal system. Once 1930 1956 1982 2008 the salt factories started to close, fluid pressures in the OGF began to recover and seawater intrusion (and mixing with hot waters) (C) TDS diminished. 25,000- 28 4. Silica and cation geothermometers 15,000 At the peak of salt production (1952-1956） the OGS was likely far from equilibrium, so that the geothermometric reser- ... voir temperatures for that period may not be correct. The following 5000 geothermometers were used: (i) Na-K-Ca (Fournier and Truesdell, 1973); (ii) cationic composition (CCG) (Nieva and Nieva, 1987); (ii) 1930 1956 1982 2008 silica (Fournier and Potter, 1982); (iv) silica (Giggenbach, 1992); (v) chalcedony (Fournier, 1992); (vi) amorphous silica (Fournier, 110- 1992); (vii) silica (Verma, 2000); and (vi) K-Mg (Giggenbach, (d) T(°℃) 1988). A summary of the general principles, assumptions, limitations 90 and applications of geothermometry to geothermal systems is given by Fournier (1981) and Giggenbach (1988). The CCG geother- mometer is an effective tool for temperature estimation (OLADE, 1994), as it utilizes four equations based on the relative propor- 70 tion of Na, K, Ca and Mg concentrations. A function of the Na/K 3 ratio is used in the case of high temperature geothermal fluids with low Mg content, as the Mg-based geothermometers reflect 50 1930 1956 1982 2008 (Truesdell, 1995). Year Reservoir temperatures inferred on the basis of geothermom- Fig. 12. Concentrations (in mg/L) of the main components in Obama hot spring etry are given in Table 1. The highest temperatures are those waters (also see Table 1) Group 1: 1936-1944; Group 2: 1945-1955; Group 3: indicated by the CCG geothermometer (Table 1; Fig. 13), whilst 1956-1959; Group 4: 1960-2005. (a) Changes in Na and Cl; (b) changes in K, Mg. the lowest are given by the amorphous silica geothermometer. Ca, HCO and SO4; (c) changes in TDS in mg/L; (d) changes of water temperatures The temperatures indicated by the Na-K-Ca geothermometer are (in°C). below those estimated by the CCG, which can be explained by the elevated Mg contents in the Obama thermal waters (Table 1). It is worth noting that the R correction factor of the Fournier and Potter Changes in calculated OGF reservoir temperatures reflect the (1979) Na-K-Ca geothermometer [i.e. R= (Mg × 100)/(Mg + Ca+K)] intensity and nature of the commercial operations in the area; par- is greater than 50 for most of the Obama spring samples, thus ticularly those related to salt production activities. Lower inferred suggesting that these Mg-rich fluids come from environments temperatures correspond to periods of high cold seawater intrusion that may be considerably cooler than what is suggested by the into the geothermal system. Na-K-Ca geothermometry results. Also and because of these high Before the salt factories were opened the average reservoir tem- Mg concentrations, the reservoir temperatures inferred by the perature inferred using the CCG (Tccg) was ~250 °C(Fig. 10). During K-Mg geothermometer are very low (5-43 °C; Table 1). the period of commercial salt extraction (from 1944 to 1959) those H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 237 ·Na-K-Ca (Fournier and Truesdell, 1973) 3000 Steam Conductive Group1 ·CCG(Nieva and Nieva,1987) (2775kJ/kg) heating · Silica (Fournier and Potter, 1982) Group2 2500- seam <1/ OGroup3 Silica (Giggenbach, 1992) piny Group4 Silica(Verma,2000) 2000 Chalcedony (Fournier, 1992) Steamloss or mixing conductive R:Rescrvoir boiling Amorphous silica (Fournier, 1992) 1500 cooling water fdi (5。) 380 2 3 dn 1000 P R475% R3 R2 Boiling spring line O 500 100℃ G G Seawater 300 0 3000 6000 9000 12,000 15,000 18,00021,000 cwater Chloride (mg/kg) 1IO 200 Fig. 14. Chloride-enthalpy plot for Obama hot spring waters showing the reservoir water and the mixing, boiling, and cooling lines. Group 1: 1936-1944; Group 2: ted 1945-1955; Group 3: 1956-1959; Group 4: 1960-2005. 100 65- 60- 55- 19351945 1955 1965 1975 1985 19952005 50- First salt factory opens Last salt factory closes Year 45- 40 Fig. 13. Change in estimated reservoir temperatures based on different geother- 35- mometers between 1936 and 2005. Group 1: 1936-1944; Group 2: 1945-1955; Group 3: 1956-1959; Group 4: 1960-2005. 30- 25 20 temperatures decreased to 2oo°C. After the factories were closed 15 the TccG gradually increased to ~220°C. The silica geothermome- 10 ters show similar trends (Fig. 12), despite being affected by silica Seav precipitation caused by cooling and boiling processes. 目 Two hot spring water samples with low K concentrations, col- 2000 4000 6000 800010,00012.000 14,00016,00018,000 20,000 lected in 1971 and 1975,give estimated reservoir temperatures that Chloride (mg/kg) are below those measured at the surface (discharge temperature) Fig. 15. Bromide-chloride plot for the hot water and seawater from the Obama (Fig. 13), which point to a predominantly dilute, meteoric origin of geothermal field. Group 1: 1936-1944; Group 2: 1945-1955; Group 3: 1956-1959; the thermal spring waters. Group 4: 1960-2005. 5. Chloride-enthalpy model and Cl-Br relation similar and lower than that for Groups 2 and 3. In other words, the chloride content and the enthalpy of the produced fluids reflect the Mixing models like the enthalpy-chloride model of Fournier importance of seawater intrusion before the start, during and after (1979) are very useful for determining chemical processes in active the end of commercial salt extraction operations. hydrothermal systems and for gaining understanding their hydrol- Fig. 15 shows a plot of bromide vs. chloride concentrations in ogy. These models were applied to 96 samples of Obama hot spring Obama hot spring and representative sea waters. The figure indi- waters in order to estimate silica and chloride contents in the reser- voir fluids, i.e. before the fluids were affected by mixing/dilution cates that during the period of salt production (Groups 2 and 3) the bromide contents in the thermal water increased, and that since and boiling processes as they ascended towards the surface. The these operations ceased in 1959 they are returning to their original reservoir fluid temperature for the four groups of hot spring groups values (pre-salt production times; Group 1). (see Section 3) was estimated using the Na-K-Ca geothermometer; the results are shown as R1 (Group 1), R2 (Group 2), R3 (Group 3), and R4 (Group 4) in Fig. 14. 6. Enthalpy-silica model Two trends can be inferred from Fig. 14, i.e. (1) steam loss and boiling, wherein the remaining fluid loses enthalpy and mass, and its Cl concentration increases, and (2) mixing with seawater, which ature of the hot water component in mixed waters and to estimate also results in a decrease in enthalpy and increase in Cl content. As geothermal reservoir temperatures (Fournier and Truesdell, 1974; indicated in Table 4, the amount of mixing for Groups 1 and 4 is Truesdell and Fournier, 1977). Table 4 Results from the chloride-enthalpy model. The percentage of the seawater component in the hot springs decreased recently (Group 4) due to the cessation of hot water extraction by the salt factories. This led to the recovering of the enthalpy of the reservoir fluid in the Obama geothermal field. Group Seawater (%) Rain (%) Chloride contents in the reservoir (mg/L) Enthalpy of the reservoir fluid (Rn, n= 1, 2,3, 4) (kJ/kg) Group 1 10 25 3450 816.5 Group2 30 15 11,400 667.0 Group 3 35 35 10,250 785.0 Group 4 20 50-70 6400 710.5 238 H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 600 Table 6 Isotopic concentrations in hot spring,wells,river and rain waters from the Obama Group 1 geothermal field (NEDO, 1988; OT, 2002). HS, hot spring; UZ, wells location of hot 500 Group 2 springs and wells is given in Fig. 3, those of the river and rain water sampling points O Group 3 in Fig. 4. 19400 Group 4 Year CI (mg/L) 8180 (%) Sample 8D (%) (mg/kg HS 1 1984 S60S -4.2 -36.1 300 HS 2 1984 4890 -4.2 -35 ica oirfluid（Group4） HS 3 1984 4960 -4 -31.9 HS 4 1984 4815 -4.3 200 33 HS 5 1984 4695 -4.1 -33.1 HS 6 1984 4630 4.5 -33.4 100 HS 7 1984 4740 -4.2 -31.9 Seawater Rainwater HS 8 1984 4590 -4.5 -32.7 HS 9 1984 4550 -4.5 34.1 500 1000 1500 2000 2500 HS 10 1984 4715 -4.5 -35.7 0 3000 HS 11 1984 4855 -5.3 Enthalpy (kJ/kg) 36.2 HS 12 1984 4840 -4.3 -35.9 HS 13 1984 4800 -4.2 -32 Fig. 16. Silica-enthalpy mixing model for the Obama hot spring waters. Ps indi- HS 14 1984 5030 -4.2 -30.8 cates the silica concentration in the reservoir fluid. Group 1: 1936-1944; Group 2: HS 15 1984 3975 4 34.7 1945-1955; Group 3: 1956-1959; Group 4: 1960-2005. HS 1 2001 4620 -4.5 34.2 HS 2 2001 4680 -4.1 -34.2 The enthalpy values used to estimate the Obama reservoir tem- HS3 2001 4410 -4.5 35.6 perature (Fig. 16) were derived from measured spring discharge HS 4 2001 4560 -4.5 34.9 temperatures and steam table data (Keenan et al., 1969). The esti- HS 5 2001 4470 4.5 35.2 HS 6 2001 4470 -4.5 -34.6 mated reservoir temperatures and silica contents are presented in HS 7 2001 4370 -4.6 -36.8 Table 5, and plotted as Ps-1-Ps-4 (for Groups 1-4, respectively) in HS 8 2001 4470 -4.5 -36 Fig. 16. Note that the temperatures are likely to be underestimated HS 9 2001 4540 -4.5 -36.5 due to the likelihood of silica precipitation caused by cooling and HS 10 2001 4650 -4.2 36.1 boiling. HS 11 2001 4710 -4.1 -35.4 HS 12 2001 4530 -4.3 -35.3 HS 13 2001 4650 -4.2 -35.2 7. Isotope chemistry, hydrothermal fluid origin, and HS 14 2001 4620 -4.3 -34.2 mixing processes HS 15 2001 4620 -4.3 -34.5 HS 16 2001 4720 4.1 33.6 HS 17 2001 4800 4.2 33.2 The deuterium and oxygen isotopic contents in thermal water are generally a good indicator of fluid origin (Craig, 1961), although UZ-15 1984 3220 -5.3 -36.2 UZ-16 1984 4830 -4.3 33.0 several processes may affect them. UZ-17 1984 4510 -4.5 -33.4 The discharge of springs located in coastal regions commonly UZ-18 1984 4590 -4.1 -33.1 UZ-19 1984 4700 4.5 35.7 tal (meteoric) water (Sveinbjornsdottir et al., 1986; Barragan et al., UZ-02 1984 4730 -4.2 -32.0 2001). The OGF has many similarities with the Reykjanes geother- UZ-20 1984 4790 -4.2 -35.0 UZ-21 1984 4910 -4.2 -36.1 mal field. Both show small surface thermal activity, and present UZ-22 1984 4830 -4.0 -34.7 the effects of seawater mixing phenomena due to their location UZ-23 1984 4700 -4.3 35.9 near the coast and the high permeability of their volcanic rocks UZ-03 1984 4730 -4.2 -30.8 (Sveinbjornsdottir et al., 1986). UZ-04 1984 4830 -4.0 -31.9 UZ-05 1984 4690 -4.2 -31.9 Data on the isotopic (18o, D) composition of waters from 15 UZ-06 1984 4510 -4.5 32.7 Obama wells and 15 Group 4 hot springs (analyzed in 1984; NEDO, UZ-07 1984 4480 -4.5 -34.1 1988), from 17 Group 4 hot springs (analyzed in 2001; OT, 2002), River water 1984 5.5 -7.4 45.9 from a river (analyzed in 1984) and local rainfall (analyzed in 1999; unpublished data, Geological Survey of Japan, 2010) are given in Rain water 1999 5.4 -7.37 -46.7 Table 6 (no isotopic data for water samples from Groups 2 and 3 exist). The purpose of these analyses was to obtain a better under- standing of the hydrological processes occurring in the OGF. As Meteoric World Meteoric Line (GMWL; Fig. 17), whose equation is shown in Fig. 17 the isotopic composition of cold waters clearly (D =8.138180+ 10.8; Rozanski et al., 1993), thus indicating these differs from that of thermal geothermal fluids. waters have a local meteoric component. Also shown in Fig. 17 is All Obama geothermal waters from the NEDO and OT stud- another meteoric line proposed by Sakai and Matsubaya (1974) for ies (i.e. those sampled in 1984 and 2001) plot close to the Global Japanese precipitation (8D = 88180+ 17). Table5 Group Estimated reservoir Silica concentration in the Fraction of hot water in the temperature determined reservoir fluid (mg/L) mixed discharge from the enthalpy-SiO2 inferred from the model(°C) enthalpy-SiO2 model Group 1 191.5 227 0.91 Group 2 171.5 165 0.67 Group 3 177.0 182 0.89 Group 4 204.0 265 0.76 H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 239 O Local groundwater Most“shifted”thermal water CAMPIFLEGREI ? OSeawater Alteredrock Oxygen-isotope shift Unalteredrock ObamaGeothermal Field(OGF RiverwateratOGF ★Rain waterat OGF 50 WAIRAKEI (0%) 8D SALTONS 100 -15 -10 5 10 (0%)OS types of waters. GMWL, Global Meteoric World Meteoric Line;lines parallel to it represent fractions of andesitic fluid that possibly contribute to the formation of the thermal waters (Giggenbach, 1992). The apparent “oxygen shift" in the Obama thermal spring and e ( na n well waters shown in Fig. 17 could be due to isotopic exchange model for the region beneath the Unzen volcano that incorpo- with the host rocks and the result of mixing of local groundwater rated the four known large geothermal systems on the Shimabara with 180-enriched magmatic water, such as the “andesitic water" Peninsula, which are (from west to east): the Obama hot springs, of Taran et al. (1989). This andesitic fluid is inferred to undergo iso- the West Unzen High Temperature Body (WUHTB), the Unzen topic exchange with the reservoir rocks, lose steam during ascent, fumarolic field and the Shimabara hot springs. Their model included and, near the surface, mix with meteoric water (ground, river and threefluid sources inferred from geodetic datacollected during the sea waters) and boil, leading to thermal spring and well waters 1990-1995 Unzen volcanic eruption. Source A is beneath the Unzen with an elevated heavy isotope content. The fraction of magmatic volcano, while B is under the Unzen fumarolic area. Source C, below andesitic fluid that contributes to the formation of the Obama ther- the WUHTB, is considered to be a magma reservoir at about 8 km mal waters is estimated at 0.08. depth. Using the mixing line with seawater shown in Fig. 17 the esti- mated proportion of seawater in Obama hot spring waters is about 9. New proposed conceptual hydro-geochemical model for 35%, assuming an end point that has the same deuterium content the Obama geothermal system as the local river water, and a level of seawater mixing consistent with the enthalpy-chloride relationships shown in Fig. 14. A schematic hydro-geochemical cross-section of the OGS is pre- sented in Fig. 19, which includes the three fluid sources in the Fujimitsu et al. (2006) model (see Section 8). The features for the 8. Earlier proposed models for the Obama geothermal eastern part of section were based on work done by Ohsawa (2006), system while the locations of the magmatic sources beneath the OGS and the Unzen volcano are from Fujimitsu et al. (2006) and Ohta (2006). Different fluid and heat sources have been proposed for the OGS. In this study, we focused mainly in developing the hydrothermal That is, lateral outflow to the west from Unzen volcano, recharge model for the Obama geothermal field (OGF). to the east from a magmatic source located beneath Chijiwa Bay, and flow from multiple geothermal systems. The temperature pro- file for the ~1400 m deep well Uz-2 (Fig. 18) shows an anomaly Measured temperature (°C) between 200 and 620 m below sea level elevation that would indi- 0 50 100 150 200 cate a lateral flow of hot fluids from the Unzen volcano (Saibi et 200 na un a uosns a suo s o “ 0 (1986) that the heat source for the Obama hot springs may be the fluids coming from that volcanic system. -200 According to Matsushima and Kohno (2006) the surface man- 400 ifestations in the OGF originate from emanations associated with a magmatic chamber located 15 km beneath Chijiwa Bay. Ohsawa qe -600 (2006) suggested that the Obama hot springs are part of a “liquid- E -800 dominated, salty water geothermal system", which is primarily as ans- e s nu a -1000 vati shallow levels, the ascending thermal waters flow laterally toward -1200 the western coastal regions of the Shimabara Peninsula, where they cool conductively. Close to the surface the hot waters mix with cold -1400 groundwater to form the boiling or hot water of the Obama thermal spring system (Ohsawa, 2006). Fig. 18. Downhole temperature log for well UZ-2 (drilled in 1984). 240 H. Saibi, S. Ehara / Geothermics 39 (2010) 228-241 WEST EAST Unzenvolcano Fugen-Dake Unzenfumarolicfield Volcanicgas>800°C(HCI) Rain Heisei-Shinzan1990-1995 .... WUHTB Shimabarahot springs Water.table Meteoricinfitratior Obama geothermal field Sea Obama springs (Chijiwa Bay) Water downflow M Boiling300°℃ （150-200°℃） S Boiling&ste Andesitic magmaticfluid Depth:15 Fig. 19. Hypothetical W-E section across the Obama geothermal feld and the underlying deep sources, showing inferred hydrological processes and reservoir characteristics. Parts of the figure are based on the works of Fujimitsu et al. (2006), Ohsawa (2006) and Ohta (2006). The Obama geothermal reservoir is a two-phase, liquid- (4) The present state of the OGS is regarded as being “in recov- dominated system. A small amount of the deep-sourced hot water ery" following cessation of salt production and thermal fluid is inferred to move near vertically towards the surface, where it extraction. mixes with considerable amounts of (1) ground waters of meteoric (5) In the future, resource utilization should be managed to avoid origin, (2) intruded seawater and (3) high temperature fluids com- overexploitation of the geothermal resource that would lead to ing from the eastern side of OGF (strong lateral flow from the Unzen seawater intrusion and lowering of spring and well tempera- volcano; Ohsawa, 2006), as is observed in the UZ-2 temperature log tures. This would require predictive reservoir modeling studies (Fig. 18). and a carefully designed well and spring monitoring program. The Obama thermal spring waters are in chemical disequilib- rium with the surrounding rocks (Fig. 12). Geothermometers give Acknowledgements different estimates for the deep reservoir temperature, with fluc- tuations reflecting variations in the measured hot spring discharge The authors would like to thank Dr. Jun Nishijima, Dr. Yasuhiro temperatures. Temperatures above 200°C could be present in the Fujimitsu and the students of the Laboratory of Geothermics, deepest parts of the OGS as indicated by the CCG geothermometer, Department of Earth Resources Engineering, Kyushu University although estimated reservoir temperatures are in the 150-200 °C for their help and assistance. The authors greatly appreciate the constructive and thoughtful comments of Drs. Greg Bignall, geothermal resource. Based on these inferred temperatures it may Patrick Dobson, Sabodh Garg and Marcelo Lippmann (members be feasible to produce electricity by tapping the OGS using con- of the journal's “Editorial Team") for their careful reading of our manuscript and numerous constructive suggestions that helped recommend drilling deep geothermal wells in the eastern part of improve this paper. We thank Ms. Katie Kovac (Schlumberger, the OGF and under Chijiwa Bay with the purpose of identifying n s n s ue s and characterizing the sources feeding the OGS and to evaluate its The authors greatly appreciate the constructive and thought- geothermal potential. ful comments of Dr. R.M. Barragan (Instituto de Investigaciones The main conclusions drawn from the present study can be sum- Eléctricas, Mexico). We also thank Dr. Shinji Ohsawa (Kyoto Uni- marized as follows: versity, Japan) and Dr. Masaya Yasuhara (AIST, Japan) for their help. The first author gratefully acknowledges the financial sup- port of the Japan Society for the Promotion of Science (JSPs). This (1) Hot spring waters of the OGS are primarily of the Na-Cl type. (2) There is a significant seawater component in the Obama hot Research). spring waters. (3) Changes in reservoir temperatures estimated using silica References and cation geothermometry reflect the effects of geothermal resource utilization. The impact of salt production shows a Barragan R.,R.M., Birkle,P.,Portugal M.,E.,Arellano G.,V.M.,Alvarez R.,J.,2001. G- history of unsustainable fluid extraction (over-use) and heat chemical survey of medium temperature geothermal resources from the Baja California Peninsula and Sonora,Mexico.Journal ofVolcanologyand Geothermal mining of the Obama geothermal resource. Research 110, 101-119. 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Derivation of Na-K-Mg-Ca hydrothermal fluids of the Reykjanes and Krafla geothermal fields, Iceland - a geoindicators. Geochimica et Cosmochimica Acta 52, 2749-2765. stable isotope study. Contributions to Mineralogy and Petrology 94, 99-109. Giggenbach, W.F, 1992. Isotopic composition of geothermal water and steam Tamanyu, S,Wood, C.P, 2003.Characterization of geothermal systems n olcan discharges. In: D'Amore, F. (Coordinator), Application of geochemistry in tectonic depressions.Japan and NewZealand.Bulletin of the Geological Survey Geothermal Reservoir Development. UNITAR/UNDP, Vial del Corso, Italy, pp. of Japan 54, 117-129 (in Japanese with English abstract). 253-273. Taran, Y.A., Pokrovsky, B.G., Dubik, I.M., 1989. Isotope composition and the origin of IHES, 2002. Analytical results of Obama hot spring waters (1886-1987). Nagasaki Water in andesitic magmas. Dokl. Akad. Nauk SSR 304, 440-443. Prefectural Institute of Health and Environmental Sciences research report, Part Truesdell, A.H.,1995.Geochemical interpretation of coastal hot spring waters from 2 of Report No. 29.Nagasaki, Japan, 272 p.(in Japanese). Indonesia. Unpublished report for the U.S. Department of Energy, Idaho, USA, Kamata, H., Kodama, K., 1993. The volcanic zone as volcano-tectonic depression 14 pp. and its formation tectonics-three tectonic events caused by subduction of the Truesdell, A.H.,Fournier, R.O., 1977. Procedure of estimating the temperature of a Philippine Sea Plate under the junction of the Southwest Japan Arc and Ryukyu Arc. Memoirs of the Geological Society of Japan 41, 129-148 (in Japanese with Journal of Research of the U.S. Geological Survey 5, 49-52. English abstract). Tsukuda,E,1993.Is the Central Kyushu really rifing north to south?Memoirs of KEEA,2005.Chemical Analysis of Marina Hot Spring Waters.Kyushu Environmental the GeologicalSociety of Japan 41, 149-161 (in Japanese with English abstract). Evaluation Association, Fukuoka, Japan, 1 p. Verma, M.P, 2000. Revised quartz solubility temperature dependence equa- Keenan, J.H., Keyes, F.G., Hill, P.G., Moore,J.G., 1969. Steam Tables: Thermodynamic tion along the water-vapor saturation curve. In: Proceedings of the 2000 PropertisofWaterincludingVapor,Liquid,andSolidPhases.JohnWiley&ons Inc., New York, NY, USA, 162 pp. 1927-1932. Matsushima, ., Kohno,Y., 2006.Relation between Chijiwa caldera and activity of Watanabe, K., 1958. Thermodynamic analysis on the heat source of Obama hot Unzen volcano. Monthly Earth 28, 97-102 (in Japanese with English abstract). springs in the vicinity of Unzen volcano district. Chigaku Zasshi. Journal of NEDO, 1988. Geothermal development research document, Unzen Western Region. Geography 67, 127-152 (in Japanese). New Energy Development Organization report 15, 1060 pp.(in Japanese). Yuhara, K., Ehara, S., Futagoishi, M., Fujimitsu, Y., 1986. Heat discharge and flow Nicholson, K., 1993. Geothermal Fluids, Chemistry and Exploration Techniques. Springer Verlag, Berlin, Germany, 263 pp.
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Saibi (2010) Temperature and chemical changes in the fluids of the Obama geothermal field.txt
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Geochemical Journal, Vol. 25, pp. 203 to 222, 1991
Geochemistry of the Nigorikawa geothermal
southwest Hokkaido, Japansystem,
YUTAKA YOSHIDA
JMC Geothermal Research and Development Co., Ltd.
72-2, Sasamori, Ukai, Takizawa-mura, Iwate-gun, Iwate-ken 020-01, Japan
(Received August 15, 1990; Accepted March 15, 1991)
The Nigorikawa geothermal system is located in the Nigorikawa caldera, southwest Hokkaido. The
basement rocks are dominantly sediments, including limestone, while Tertiary rocks are mainly
andesitic volcanics. Geothermal waters in this area can be divided into four groups on the basis of their
relative C1 and SO4 contents (i.e., hot springs outside the Basin, hot springs inside the Basin, hot water
derived from faults in the pre-Tertiary Kamiiso Group, and hot water derived from the fracture zone
related to the caldera wall). Although the SO4 concentration of the thermal waters is' controlled by
anhydrite solubility in the deep formations, there are no trends indicating its dissolution. In the shallow
hot spring waters, S04 concentration decreases by mixing with groundwater. Isotopic data suggest that
the geothermal water is formed by simple mixing of meteoric water with deep hot water having a large
magmatic component and/or altered sea water. Relative He, Ar and N2 contents of Nigorikawa geother
mal fluids indicate that they are mixtures of magmatic-derived gas and atmospheric air dissolved in
groundwater. Based on trends in the Cl-enthalpy relationship, two endmembers have been identified in
the Nigorikawa system, i.e., original deep fluid and zero chloride, shallow steam-heated water; a third
endmember is now present due to the reinjection of high chloride waters from the geothermal power
development. The reservoir liquid was saturated with calcite, resulting in CaCO3 scale deposition during
production. However, after a prevention system using a CaCO3 scale inhibitor was completed in 1985,
the CaCO3 scaling problem was solved. Stevensite scale has recently precipitated in the two-phase
pipeline and waste water pipelines related to well ND-1, due to an incursion of Mg-rich low enthalpy
water.
INTRODUCTION
The Nigorikawa geothermal system is located
in the Nigorikawa Basin of southwest of Hok
kaido, Japan (Fig. 1). There are about 60 hot spr
ings in the Nigorikawa Basin, which are mainly
used for greenhouses and hotels. Since 1972,
geological, geochemical and geophysical explora
tion aimed at geothermal development has been
conducted by Japan Metals and Chemicals Co.,
Ltd. (JMC). Geothermal development began in
1977 by Dohnan Geothermal Energy Co., Ltd.
(DGE) with the cooperation of JMC and Hok
kaido Electric Power Co., Ltd. Construction of
the Mori geothermal power plant started in
1981, and operation began in November 1982.
Nine exploration wells, six production wellsand eleven reinjection wells were completed
prior to 1982. The vertical depth of these wells
ranges from 430 m to 2,360 m. Three additional
production wells and two reinjection wells were
drilled up to July 1990.
The chemical compositions of hot springs in
and around the Nigorikawa Basin were discussed
by Uzumasa et al. (1959), Matsubaya et al.
(1978) and Yoshida (1982, 1983, 1990). Isotopic
and chemical study of gases discharged from
fumaroles and wells was conducted by Nagao et
al. (1979) and Yoshida (1986). Fukutomi et al.
(1963) and Urakami and Nishida (1977) studied
the Nigorikawa thermal area from a geophysical
standpoint, and also mentioned the chemical
composition and flow mechanism of the hot
spring water.
203
204 Y. Yoshida
141E
N
43N
140E
Lake Toya
I (D O Sapporo
Lake Shikotsu
0GYMNigorikawa geothermal system
00 ~0~k
MoriKomagatake
HakodateSea of
Japan
Pacific
Ocean
50km
Fig. 1. Locality map of the Nigorikawa geothermal system. • : Hot springs outside the Basin.
The thermal waters of the Nigorikawa hot
springs are high in HCO3, and their Cl concentra
tions vary over a wide range. Compared with
thermal water of hot springs outside the Basin
and fluids discharged from wells, the Nigori
kawa hot spring waters are depleted in S04
Travertine (CaCO3) often deposits from thermal
waters of the Nigorikawa hot springs. Shortly
after beginning operation of the power plant,
production wells were plugged by CaCO3 scale
deposition. JMC and DGE investigated techni
ques for the removal of CaCO3 scale, and devel
oped a prevention system involving the direct in
jection of a scale inhibitor into the hot water at
depth before flashing.
Since operation of the power plant began,
the chemical characteristics of the discharged
fluid changed over a period of about eight years.
In this paper, previous studies on the fluid
chemistry by the present author are summarized,
and then the variation of fluid chemistry and
rare gas compositions of the geothermal fluid are
discussed, along with problems related to scal
ing.GEOLOGICAL SETTING
The geological map of the Nigorikawa Basin
and surrounding area is shown in Fig. 2. The
Nigorikawa geothermal system is located in the
Nigorikawa Basin, about 5 km from the
seacoast. The Basin is a pentagonal-shaped plain
about 3 km wide, has an elevation of 100 m, and
is surrounded by mountains rising to about 400
m. The Nigorikawa Basin is similar to a Crater
Lake type caldera, and formed during volcanism
at 12,900±270 B.P. (Sato, 1969). Dips of the
caldera wall steepen at depth from 45' to 70°,
and its diameter is less than 500 m at 2,000 m
depth (Ide, 1982). Fig. 3 shows the contour of
the caldera wall with main feed points of geother
mal wells produced in the early stage of the
development.
The rock formations around the Nigorikawa
Basin are composed of the pre-Tertiary Kamiiso
Group (slate, limestone, chert, tuff), Ebiyagawa
Formation and Katsuragawa Formation of
Neogene age (sandstone, conglomerate,
andesite, tuff), and caldera fill sediments and
Nigorikawa geothermal system 205
ae^C
v
vD
vve
=YUa Uchiura Bay
D
:•DpD.p v
4 D
Q p
A
Nigor.ikawV
0
aD
D
4Basia ti}
4
Atv
v v•+••o'1(A%A AA
AAAAA
A A A A A
A • A A
4
Q4.V4 V ''• Mt. Kenashi
v°°44apD Vr~.
•'w
K
1w 4
vv D
4 D
v''
kmv 1~"4 Dv '
Mt. Gushindake a 4 v
vV
°Dv .: p 4
'~4Q
D1Dv v.
.oD
v°a
0 1 2 3 4 5
D,.
Fig. 2.Legend
1 4 -\\]
2 A An
3
4 F471
5 6
7
8 vvv
9 R'
10 -~ u
11-f
12 -fIshikura F.
Kenashiyama lava flow
Setana F.
Gushindake vol. breccia
Kuromatsunai F.
Yakumo F.
Kunnui F. (Ebiyagava formation
Andesite dyke
Fossil
Fault
Anticline
Syncline
Geological map around the Nigorikawa Basin (after Ide and Doi, 1982).
N
~SlS• NT-306(-350
NB-2(-800) '
Z 0
D-6
NF-7(-850) •
0l\M -t
NC-1(-1200
NT-301(_BOZ11
J I 1 _750)
N(-1550) \
ND-5(-500)
NT-302(-1100)
;1500m NP-2(-900)
x-1000 m
z -500mcross section
Rim of a plane of the Basin
(100-150m A.S.L.)
Caldera margin
1l m
Fig. 3. Structual contour map of the caldera wall and distribution (location and elevation) of main feed points
of wells (at the early stage of their operation). ( ): main feed point and approximate elevation of wells near
Mt. Bozu; • ( ): main feed point and approximate elevation of wells of the pre-Tertiary fracture group; 0
( ):main feed point and approximate elevation of wells of caldera wall fracture group. Depth: in meter A.S.L.
206 Y. Yoshida
W Nigorikawa BasinE
Lake sedimentsupper limit of
anhydrite distribution
TertiaryCaldera fill
deposits
pre-Tertiary10.b
t
10V
"VV
'VVV
V V
VVV
VVNF-2/
NT-302Tertiary
pre-Tertiary
500m0
-500
-1000
-1500
-2000
-2500(in ma.s.l.)
Fig. 4. Schematic cross section of the Nigorikawa Basin. 0: main feed point of well of caldera wall fracture
group. Main feed points of ND-3, ND-7 and NF-1 exist in the pre-Tertiary formation at depths from -1950 to -2300 m A.S.L.
Ishikura Formation (silt, tuff, volcanic ash) of
Quaternary age. The pre-Tertiary formation is
unconformably overlain by Neogene forma
tions. An andesite dyke intruded into the caldera
fill deposits, and is inferred to be the root of the
central dome lava of the Nigorikawa caldera.
The upper portion of the caldera fill is composed
of lake sediments. A schematic cross section of
the Nigorikawa Basin is shown in Fig. 4.
One of the northeast-southwest transverse
faults in the Kamiiso Group is an excellent pro
duction zone when intersected by a geothermal
well. The fractured zone related to the caldera
margin is also one of the productive zones of this
geothermal system.
Geothermal activity in the Nigorikawa area
includes hot springs in and around the
Nigorikawa Basin and Mt. Komagatake, which
is an active volcano 20 km east of the Basin.
Cold gas discharges to high temperature
fumaroles (up to the boiling point) are located in
the Basin; their gas components are mainly C02
and H2S. Alunite, kaolinite, montmorillonite
and amorphous silica alteration of the
Ebiyagawa Formation is present at the northeastN
lowAltered
Z o n e
.old gas discharge
chloride _'°( T
area
,i619 ' PMt.
400w deep
contour ofUchiura
Bay
Bozu / at
4
3.
a'•..2a
(-300w A.S.L.)
caldera wallN i g o r i k away
Basin
1 km
Fig. 5. Distribution of hot springs, cold gas
discharges and altered zone of the Nigorikawa Basin.
0: High chloride; e: Medium chloride; •: Low
chloride.
end of the Nigorikawa Basin (Igarashi et al.,
1978). The distribution of hot springs inside the
Nigorikawa geothermal system 207
Table 1. Chemical compositions of some thermal waters of the Nigorikawa hot springs (modified from
Yoshida, 1982). Thermal waters are divided into three groups, with Cl concentrations less than 100 mgl1, 100
1000 mg / 1, and greater than 1000 mg / l
Temp.
No. (°C)pH HCO3 CO (mg/1) (mg 71) Ci
(mg/1)S04
(mg/1) Na
(mg/1)K
(mg/1) Ca
(mg/1)Mg
(mg/1)Si0
(mg /i)
Low
chloride
Medium
chloride
High
chloride1)
2)
3)
4)
5)
6)
7)
8)
9)
10)
11)
12)
13)
14)
15)
16)
17)
18)
19)
20)
21)40.3
51.2
54.3
57.3
58.5
57.4
60.7
50.3
65.8
46.8
42.1
76.2
61.4
73.5
88.5
75.6
76.6
67.5
56.6
72.7
55.67.49
6.28
6.81
6.38
7.90
6.53
6.31
6.26
6.48
6.15
6.14
6.57
6.22
6.61
6.88
6.45
7.11
6.63
6.51
6.51
6.75144
351
290
583
311
1050
452
633
883
621
696
801
444
760
927
668
847
637
460
460
11208
309
75
408
7
519
371
584
490
738
845
362
449
313
205
397
110
245
238
238
3359
12
13
19
23
43
78
163
240
476
521
645
860
909
1750
1870
2100
2330
2600
2700
36905
8
10
19
10
33
8
12
13
57
70
10
7
73
21
26
23
25
26
31
7621
53
61
89
61
183
87
187
248
420
448
583
555
801
1110
1220
1180
1480
1630
1670
226017
34
35
47
32
31
19
26
102
23
11
28
30
9
21
57
63
34
35
34
10214
34
13
73
13
169
110
108
127
14
109
117
94
74
109
137
229
120
85
96
1048
36
6
42
6
36
28
73
60
54
52
19
60
44
87
47
162
50
42
61
184101
149
161
153
177
115
175
143
207
175
162
188
205
128
207
133
255
181
184
178
221
(PH: measured at room temperature)
Basin, and the area of cold gas discharges and
the altered zone, are shown in Fig. 5. Anhydrite,
calcite, montmorillonite, pyrite and sericite are
found in drillcores and cuttings. However,
anhydrite does not occur in the shallow portion
of the caldera fill (at less than 200 m depth; Fig.
4).
FLUID CHEMISTRY
Fluid composition before exploitation
The chemical compositions of the
Nigorikawa hot springs are given in Table 1. The
thermal water of the Nigorikawa hot springs is
characterized by a wide range in Cl concentra
tions (9 to 3,690 mg/1) as compared with concen
trations of HCO3 + C02 (152 to 1,570 mg/1) and
Si02' (101 to 255 mg/1) (Yoshida, 1982, 1990).Hot springs are unevenly distributed in the nor
thern half of the Basin (Fig. 5). Furthermore, the
location of hot springs containing less than 100
mg Cl/l is limited to the central and northwest
part of the Basin.
Compared with the thermal water of hot
springs outside the Nigorikawa Basin and hot
water discharged from wells in the Basin, the
Nigorikawa hot spring waters are very low in
SO4 (Yoshida, 1982, 1990). Such character comes
from the chemical and physical processes occur
ring in the Nigorikawa hot spring system, which
include steam separation, anhydrite deposition
and mixing with surface water as the deep waters
rise to the shallow reservoir feeding the
Nigorikawa hot springs.
Based on the relationship between Cl concen
tration and enthalpy (from the quartz geother
208 Y. Yoshida
v300
200
100
Surface water•A
• ••B Deep parent fluid (1985-1986)
Deep parent fluid (1977-1981)X
0
2000 4000 6000
a (mg/kg)
Fig. 6. Relationship between Cl concentration and enthalpy for hot springs (after Yoshida, 1982), based on
quartz geothermometer temperature. •: hot spring; O: average value of each production test period for wells
of the pre-Tertiary fracture group; O: average value of each production test period for wells of caldera wall frac
ture group. Lines A and B are regression lines for waters from wells of pre-Tertiary fracture group (group (1) in
text) and caldera wall fracture group (group (2) in text), respectively. Regression lines for two well groups were
obtained by the least squares best fit method for the data of production tests carried out from 1977 to 1981. Line
A: h=0.0129 x C1+202 (n=85, r=0.533). Line B: h=0.0212 x C1+151 (n=44, r=0.836).
mometer) of the Nigorikawa hot spring waters
(Fig. 6), Yoshida (1982) suggested that the hot
spring water has as its source hot water ascen
ding through the fractured zone along the
caldera wall, with dilution by a low chloride end
member with a temperature of about 150°C (i.e.
a steam-heated water).
Discharge testing of wells was conducted in
the early stage of the assessment and develop
ment. Based on the relationship between Cl con
centration and enthalpy (from the quartz
geothermometer), the S04/Cl ratio and the loca
tion of the feed point of each well, two groups of
hot water were recognized (Yoshida, 1981, 1982):
(1) hot water derived from faults of the
Kamiiso Group,
(2) hot water derived from the fracture zone
along the caldera wall.
Because of the low productivity of wells drilled in the northeast portion of the Basin (near
Mt. Bozu), four wells. (NT-301, NT-306, NB-2
and NC-1) are excluded from this discussion.
Regression lines for waters discharged from
wells of each group were determined from the
relationship between CI concentration and en
thalpy of hot waters in the reservoir before
flashing (Fig. 6). Regression equations, number
of data and their correlation coefficients are also
shown in the figure. Although the reservoir com
positions deduced for the Nigorikawa hot spring
waters are scattered, they tend to fit an extension
of the regression line of group (2) described
above. Therefore, hot water ascending along the
caldera wall fracture zone is regarded as the
source of the hot spring waters. The data trends
in Fig. 6 also indicate that low chloride, low en
thalpy water does not interact with thermal
waters in the Nigorikawa Basin. This implies
Nigorikawa geothermal system 209
hiah
55.3
763HOT SPRING
,-~li,,it„ low salinity
2361
10.7x 1057.4
31.9759
3.6x 101 10.7 x 101) SURFACE WATER
0m
RESERVOIR
171.4 236110
.02755
0
24.8x10
763 38.1x101164.2
31.9759
11.7X107
150m
201 1321
568016.1 x 101
HOT WATER666
044
2.9x10
TEAM
400-1500m 257
5000365
9.0 x 101
DEEP HOT WATER16.0x101
FLUID
flow rate kgfnin.
heat flow Win.CONDUCTIVE
heatftow cal/mi
Fig. 7. Summarized heat and fluid balance in
Nigorikawa Basin (after Yoshida, 1982).that shallow cold ground water does not mix
with the thermal water; rather, the latter results
from various mixtures of deep chloride fluid and
shallow steam-heated water.
Due to the wide variation of Cl concentration
in Nigorikawa thermal waters, Yoshida (1982)
deduced that boiling of the deep fluid must occur
during its ascent, with subsequent formation of
a steam-heated water. That is, the enthalpy of
the deep parent fluid feeding the Nigorikawa hot
springs has a minimum temperature of about
250°C (about 260 kcal/kg), typical for the fluid
of the caldera wall group (on the regression line
of group (2)). This fluid begins to boil during as
cent at a depth of about 400 m. This depth is
based on two reasons; 1) the vapor saturation
pressure of 250°C water is about 40 bars, and 2)
the the low chloride hot springs are located over the
400m deep (-300m A.S.L.) contour of the
1000
so4 (mg/1)
500
00 4
b
•/ &
0/
o
000wtio
G1
110111, 90~/
o~
--b
09 (D
' NT -302 m ND-5
m
NF-2
(1979.4.21)- S04/C1-0.03
m NF-7
2000 4000 6000 8000
C1 (mgn)
Fig. 8. Relationship between concentrations of Cl and S04 (hot springs: after Yoshida, 1990). • : Hot spring
compositions outside the Basin; o : hot spring compositions inside the Basin (Nigorikawa hot spring); O: water
discharged from geothermal wells of the pre Tertiary fracture group; O: water discharged from geothermal
wells of caldera wall fracture group.
210 Y. Yoshida
caldera wall (Fig. 5). The liquid phase ascends
along the caldera wall while the steam rises ver
tically. The low chloride versus medium to high
chloride hot springs are separated, with the low
chloride and high chloride hot springs located
near the central portion and north end of the
Basin, respectively. This is consistent with the
model of boiling from 400 m depth as waters
ascend along the caldera wall fracture. A sum
mary heat and fluid transfer model is shown in
Fig. 7. This model was considered on the basis of
measured flow rate, Cl concentration and temper
ature of each hot spring water, reservoir tempera
ture based on the silica geothermometer, and
conductive heat loss (Fukutomi et al., 1963;
Urakami and Nishida, 1977) from the ground
surface.
The variations in S04 concentrations in the
hot spring waters in and around the Nigorikawa
Basin and in the well discharges are related to the
solubility of anhydrite and to mixing with dilute
waters. As shown Fig. 8, geothermal waters of
this area are classified into four groups based on
their S04/ CI ratio. Compared with the SO4 / Cl
ratio of sea water, the ratio of hot spring water
outside the Basin (Fig. 1) is high, while that of
hot spring water inside the Basin is very low. The
sulfur isotopic compositions of SO4 in thermal
water of hot springs outside the Basin
(S34S = +20.8 to +27.4%o) indicate that SO4 is
derived from anhydrite in the Tertiary
Ebiyagawa Formation. Yoshida (1983) conclud
ed that variation of the SO4 concentration in the
Nigorikawa hot spring waters was due to the
following process.
Circulating groundwater, steam-heated and
charged with CO2 gas from deeper levels, derives
its Ca and Mg by water-rock interaction. This
water then mixes with and dilutes the ascending
deep hot water, which is high in CI and S04; the
Ca and SO4 are then eliminated from the mixture
by deposition of anhydrite in the high tempera
ture reservoir, according to the following reac
tions:
CaCO3+CO2+H2O Ca2++2HCO3 ,
Ca2+ + SO4 -a CaS04J..S'Bo ('6 SMOW)
-10 -50 •,5
-1
hSD
-20
O ~
0
O®NT--301
--40 ND-6 ®NT-306
0 -5 NT-302
0
00
i 0 --60
(%.SMOW)
Fig. 9. Relationship between aD and S18O(after
Yoshida, 1990). • : Hot spring compositions outside
the Basin; 0: Hot spring compositions inside the
Basin (Nigorikawa hot spring); O: water discharged
from geothermal wells of the pre-Tertiary fracture
group; O:water discharged from geothermal wells of
caldera wall fracture group; x : river water.
-20000
Sea Wa
-15000
CI
10000
mg/I
Local Meteoric Water
5000
0
i
80-60
SO-40 ,h SMOW-20 0
Fig. 10. Relationship between 6D and Cl concentra
tion (after Yoshida, 1990). The symbols are the same
as those in Fig. 9.
Further dilution of this fluid occurs at shallow
depths, hence the SO4 concentration of the
geothermal fluid becomes extremely low in hot
spring discharges, and distribution of anhydrite
in wellbores is limited to greater than about 200
m depth (Fig. 4).
In spite of the change in chemical characteris
tics, the relationship between 8D and 8180 of
waters indicates a simple mixing of deep fluid
with thermal water of the Nigorikawa hot
Nigorikawa geothermal system 211
Table 2. Chemical compositions of hot waters discharged from wells in the initial stage of exploitation
Well Sampling date Tqtz pH T-C02 Cl SO4
(°C) (mg/1) (mg/1) (mg/1) Na
(mg/1)K
(mg/1) Ca
(mg/1)Mg
(mg/1)SiO
(mg /1) Fe
(mg/1) Al
(mg/1)H2S
(mg/1)
NT-301
NT-302
NT-306
NB-2
NC-1
ND-1
ND-3
ND-5
ND-6
NF-1
NF-2
NF-2
NF-7
NF-91977. 3.14
1978.11.15
1978. 6. 5
1979. 3. 1
1979. 5.29
1979. 4. 4
1979. 1.25
1979. 4.25
1978.11.25
1981. 3.31
1979. 4.21
1980. 3. 2
1979. 7.21
1980. 7.13211 8.70 493
237 8.70 107
224 8.82 125
226 8.30 251
219 8.76 429
249 8.21 273
257 8.70 352
238 8.24 169
259 7.97 173
262 9.06 275
237 8.73 75
235 8.72 176
253 8.51 48
245 9.25 3145470
5820
5550
5250
5800
7070
6790
6400
5520
6990
5790
4780
7370
5880382
124
232
243
429
367
358
189
651
279
165
188
113
5972840
3580
3420
3610
4120
4710
4220
4150
3750
4200
3720
3090
4440
4230441
554
487
593
504
749
692
720
361
812
686
496
809
470 5.0
22.3
20.9
14.0
6.4
7.8
5.6
15.2
36.9
5.7
17.3
10.6
25.8
5.72.2
3.1
5.4
1.1
2.6
0.3
0.4
1.5
0.1
0.4
2.0
0.7
1.7
0.3399
579
481
497
446
677
748
583
768
801
576
559
710
6420.3
0.4
0.3
0.5
0.4
0.2
0.4
0.1
0.4
0.4
1.7
0.1
0.7
0.20.1
0.4
0.7
0.7
0.6
<0.1
0.5
0.2
0.8
1.0
0.8
0.1
0.8
0.42.7
1.7
9.3
1.7
2.7
1.4
9.3
<0.1
<0.1
7.1
1.8
2.0
2.4
6.7
Collected at atmospheric pressure
Tqtz: temperature of quartz geothermometer (adiabatic cooling)
pH. measured at room temperature
T-C02: (H2CO3 + HCO3 +C02-) as C02
Table 3. Chemical compositions of steam discharged from wells in the initial stage of exploitation
Well Sampling datePressure of
separator
(kg/ cm2G)Steam fraction
in discharge
at pressure(mole %)
Total gas
in steamCO2 H2S(mole %)
H2 N2 CH4
NT-301
NT-302
NT-306
NB-2
NC-1
ND-1
ND-3
ND-5
ND-6
NF-1
NF-2
NF-2
NF-7
NF-91977. 3.14
1978.11.15
1978. 6. 5
1979. 3. 1
1979. 5.29
1979. 4. 4
1979. 1.25
1979. 4.25
1978.11.25
1981. 3.31
1979. 4.21
1980. 3. 2
1979. 7.20
1980. 7.134.0
5.9
6.0
6.0
5.5
5.9
6.0
7.8
5.9
5.3
5.9
6.0
4.7
6.00.13
0.16
0.13
0.16
0.12
0.20
0.21
0.14
0.22
0.23
0.16
0.15
0.21
0.186.9
5.5
3.7
5.8
7.9
3.7
5.4
2.6
2.4
5.3
5.6
2.9
4.1
3.598.3
97.5
98.1
97.6
98.8
97.9
97.9
97.4
97.5
97.7
97.6
97.4
98.0
98.21.1
1.6
1.6
1.7
0.7
1.0
1.0
2.3
1.2
1.1
1.6
1.5
0.9
0.50.017
0.005
0.017
0.046
0.004
0.023
0.007
0.001
0.079
0.010
0.003
0.008
0.020
0.0030.371
0.481
0.163
0.398
0.318
0.538
0.471
0.173
0.679
0.407
0.453
0.729
0.614
0.6810.212
0.415
0.119
0.256
0.179
0.539
0.623
0.117
0.542
0.593
0.344
0.363
0.466
0.616
springs (Yoshida, 1990) (Figs. 9 and 10),
most of the dilute springs similar to
meteoric water in isotopic composition.with
local
Geothermal fluid in the initial stage of exploita
tion
After liquid-vapor separation at the surface,
the waste water is reinjected. This has resulted in
a gradual change in the temperature and
chemical characteristics of the reservoir fluidfrom its natural state. The chemical characteris
tics of the Nigorikawa geothermal fluid prior to
exploitation may be determined from produc
tion test data, obtained before the start of power
plant operation.
Samples were collected from wells from 1977
to 1981, shortly after their initial discharge.
Steam samples were collected from a steam line
after liquid-vapor separation, while water
samples were collected at atmospheric pressure
212 Y. Yoshida
from a sampling valve attached to a two-phase
line near the well head. The chemical composi
tion of the hot water and steam for each well is
shown in Tables 2 and 3, respectively.
The concentration of non-condensable gas in
steam (at the separation pressure) ranges from
2.4 to 7.9 mole%, higher than that of other
geothermal power plants which are now in opera
tion in Japan. CO2 concentration ranges from
97.4 to 98.8 mole% of the non-condensable
gases, and from 0.4 to 0.8 mole% of the total
discharges.
The hot water is weakly alkaline after
flashing, and has a range of Cl concentration
from 4,780 to 7,370 mg/1, sampled at at
mospheric pressure; this corrects to 3,520 to
5,140 mg / kg at reservoir conditions. The SO4
concentration ranges from 113 to 651 mg/i (79
to 446 mg/kg in the reservoir), higher than that
of the Nigorikawa hot springs. As mentioned
before, well waters can be classified into two
groups. One is derived from a fracture zone in
the pre-Tertiary formation, the other is from the
caldera wall fracture zone. Based on the low
SO4/Cl ratio of hot water, less than 0.03, waters
from NT-302, ND-5, NF-2 (1979.4.21) and NF-7
correspond to the latter group (Fig. 8). Wells
located near Mt. Bozu, NT-301, NT-306, NB-2
and NC-1, produce from fractures in the pre-Ter
tiary formation outside the caldera, though they
are poor producers; wells NB-2 and NC-1 are
now used as reinjection wells. Wells ND-l, ND
3, ND-6, NF-1, NF-2 (1980.3.2) and NF-9 also
produce from the pre-Tertiary fracture group.
Many chemical data were obtained from pro
duction tests carried out in the early stage of
development. Because of the variation of en
thalpy and Cl concentration of hot water from
each well, a statistical analysis was done. The
characteristics of the endmember fluid in the mix
ing series may be determined from chloride-en
thalpy relations (Fig. 6). Two trends originate
from the same high chloride, high temperature
parent (nearly 6,000 mg/kg Cl at about 270°C).
These two dilution trends project to zero
chloride at different enthalpies, roughly at 150
and 200°C. Steam-heated water of zero chloride,6000
CI
4000
mg/1
2000
0w~ m
/ 0
Br10
Mg/120
Fig. 11. Relationship between Cl and Br.
bols are the same as those in Fig. 9.The sym
equilibrated at about 150°C, discharges from the
Nigorikawa hot springs (Yoshida, 1982). Water
of zero chloride and equilibrated at a higher tem
perature (up to 200°C) is caused by a greater
degree of steam heating at depth.
The range of deep equilibration temperature
of the steam-heated diluent, and therefore of the
mixtures of ascending water, suggests the ex
istence of different flow paths of surface waters.
Surface water is steam-heated to a greater degree
along the path related to the pre-Tertiary frac
ture zone, in comparison to the caldera fracture
zone. The distinct SO4/ Cl ratios of geothermal
fluid in these two environments relate to the
different paths of surface water flow.
Figure 8 shows the relationship between con
centration of Cl and S04 of hot spring waters
and hot water discharged from wells. Like the
hot springs outside the Nigorikawa Basin, water
discharged from wells drilled outside the caldera
(into the pre-Tertiary fracture zone) is con
siderably enriched in SO4. In contrast, water
discharged from wells inside the caldera has been
diluted by shallow water which has lost S04 near
the surface.
The Cl, SO4 and Ca concentrations in the hot
spring water and water discharged from wells are
variable, while other components are not distinc
tive. The Br/Cl ratios of hot spring water and
water discharged from wells are similar to that
Nigorikawa geothermal. system 213
of sea water (Fig. 11).
The relationship between 8D and 8180 in
dicates a simple mixing of deep fluid having
heavy 8D and 8180 with local meteoric water
(Fig. 9); Fig. 10 also indicates a simple linear 8D
Cl relationship. These data suggest that the
geothermal fluid of this area is formed by the
mixing of surface water with a high temperature
fluid. Concerning the high temperature fluid, the
following two origins are possible.
(1) Altered sea water. A mixture of sea water
and meteoric water modified its isotopic composi
tion and Cl concentration through hydrogen
isotope exchange between the fluid and clay
minerals and/or by steam separation in a high
temperature reservoir. Additional oxygen
isotope exchange occurred between fluid and car
bonate and silicate minerals in the reservoir.
(2) A volcanic origin has been proposed for
the isotopically heavy waters, as noted in other
Japanese hydrothermal systems (Mizutani, 1978;
Mizutani et al., 1986).CHANGES OF CI CONCENTRATION AND
ENTHALPY WITH TIME
Since the Mori geothermal power plant was
put into operation in 1982, the Cl concentration
and enthalpy of deep waters supplying produc
tion wells have changed. Figures 12 to 15 show
the changes with time in Cl concentration and en
thlpy of ND-1, ND-5, ND-7 and NF-1. The Cl
concentration of the reservoir liquid is obtained
by correcting for steam loss to the quartz geother
mometer temperature. ND-7, first discharged in
1985, is a production well drilled into the same
pre-Tertiary fracture zone as NF-1, and is also
an excellent producing well (feed points: -2200
to -2300 m A.S.L.).
Though the trends with time for Cl concentra
tions and enthalpies for various production wells
show complicated patterns, several important
trends can be recognized.
(1) Increasing Cl concentration and decreas
ing enthalpy for a short period after operation
(ND-1).
(2) Increasing Cl concentration and enthalpy
300
290
280
o, 270
m 260
x
250
n
240
s
w 230
220
210
200ND-1
O
ft: Enthalpy
. C1
31E
* '
0
0
000
o e'0
0 0
o 0! o 0°
0 00
0~0
Aeon:
00 00000
04%
9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9
184 185 186 187 188 189 190
Changes with time in Cl concentration and enthalpy of ND-1.12
I10000
9000
8000
7000
6000
5000
4000
3000
2000
1000
9 12 3 6
83
Fig. 12.0E
a
a
c
0
m
c as
c
0
c~
214
300
290
280
270
260
250
240
230
220
210
200Y. Yoshida
m
T
GL
m
r
r
wND-5
O : Enthalpy
:a
O*~Oy~
O0 0 % e~. ~O'.
%*Os~ me
O
OO
912 3 6
183
Fig. 13.9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12
184 185 186 1 87 188 189 190 1
Changes with time in Cl concentration and enthalpy of ND-5.10000
9000
8000
7000
6000
5000
4000
3000
2000
1000
0E
0.
CL
c
0
d
V
C
V
C-)
m
0
T
CL
t
c
W300
290
280
270
260
250
24.0
230
220
210
200ND-7
O : Enthalpy
:Cl
A
O
O X00
000
s f%
09 12 3 6
183
Fig. 14.9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 912 3 6 9 12 3 6 9 12
184 185 186 187 188 1 89 190 1
Changes with time in Cl concentration and enthalpy of ND-7.10000
9000
8000
7000
6000
5000
4000
3000
2000
1000
0E
o.
0.
c
0
w
L
C
d
V
C
0
0
U
Nigorikawa geothermal system 215
300
290
280
270
260
`r 250
240 m
s w 230
220
210
200NF-1
p : Enthalpy
:C1 0 O
-o
0 o
* 0
001*
~ W WO
mow
a >O
O
00 ep cb
M ON 0ep OQP
O
912 3 6 9
183
Fig. 15.12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12
184 185 186 187 188 189 190 1
Changes with time in Cl concentration and enthalpy of NF-1.10000
9000
8000
7000
6000
5000
4000
3000
2000
1000
0E
n
G.
C
0
w
c c
v
for a short period after operation (NF-1).
(3) Constant Cl concentration and enthalpy
for three years after operation, with both
decreasing after start of operation of ND-7 (ND
5).
(4) Increasing Cl concentration and decreas
ing enthalpy after operation of ND-7 (ND-1,
ND-7 and NF-1).
(5) Decreasing Cl concentration and en
thalpy after the middle of 1986 (ND-1, ND-7 and
NF-1).
These observations are interpreted as in
dicating the following possibilities, depending
on the situation.
(1) Mixing of ascending deep fluid with rein
jected waste water (high chloride, low enthalpy).
(2) Upflow of deep fluid is accelerated by a
decrease of the reservoir pressure due to produc
tion.
(3) Low chloride and low enthalpy surface
water is drawn into the reservoir following
pressure decrease.
These three possibilities can be illustrated in
Fig. 16. Most Cl concentration and enthalpy
data for the reservoir liquid plot within a trianglewhose apices are the three endmembers of:
(1) Deep parent fluid (C1= 6, 500 mg / kg, en
thalpy = 300 kcal / kg). Values of Cl concentra
tion and enthalpy are predicted from the intersec
tion of trends of compositions of reservoir
waters from ND-5, ND-7 and NF-1. These
values are also similar to those obtained from
the cross point (C1= 6,150 mg / kg, en
thalpy=281 kcal/kg) of the two regression lines
of Fig. 6 (Yoshida, 1982). The slight difference
between these two parent composisions may be
due to the different period of production. The
former was obtained from data of 1985-1986,
the latter from those of 1977-1981 (suggesting a
hotter, more chloride-rich fluid being drawn into
the deep system due to production).
(2) Zero chloride hot water (C1= 0 mg / kg,
enthalpy=170 kcal/kg). This point is obtained
by extrapolation of trend lines of ND-5 and NF
1 toward zero chloride. The enthalpy is nearly
equal to the mean value (C1= 0 mg / kg, en
thalpy=176 kcal/kg) of the zero chloride hot
waters of Yoshida (1982) (Fig. 6).
(3) Reinj ected waste water (C1=
7,000 9, 000 mg / kg, enthalpy =120 kcal / kg).
216 Y. Yoshida
350
r~o
250
.~d
a
200
r~
W
150
liiD:5 Ot0
~4~/1P IDeep parent fluid
Nigorikawa
hot springZero chloride
hot waterr c nt trend ND-7 and NF-1
(after ND-7 operation)
Reinjected waste water
Fig. 16.
system.0 2000 4000 6000 8000 10000
Cl concentration (mg/kg)
Relationship between Cl concentration and enthalpy of reservoir fluid of the Nigorikawa geothermal
• : Cross point of two regression lines of Fig. 6 (after Yoshida, 1982).
Although the Cl concentration is variable, its en
thalpy is fixed.
Because of the mixing of three endmembers
in various proportion, changes with time are not
simple. However, the individual values of Cl con
centration and enthalpy are confined within the
envelope. In some exceptional cases, simple
single component trends are present.
In the case of ND-1 just after start of its
operation, monthly data defined a trend to rein
jected waste water. After discharging ND-7, the
monthly compositions of water from ND-1, ND
7 and NF-1 also migrated toward the injected
fluid composition. This indicates the return of
reinjected waste water into the production zone.
In the early stage of production, the composition
of NF-1 water changed toward the predicted
deep parent value, due to an increased flow of
deep fluid. In the period when both Cl concentra
tion and enthalpy decreased for ND-5 and NF-1,
the trend was toward the zero chloride hot
water. This trend suggests that steam-heated
water started to be drawn into the reservoir from
a shallow horizon. No trend to cold groundwaterwas observed, indicating an extensive "cap" of
the 150°C steam-heated waters over the upflow.
As described above, mixing of three kinds of
water, i.e. original fluid from a deep horizon,
zero chloride hot water derived from steam
heating at shallow levels, and waste hot water,
now occurs in the geothermal reservoir of the
Nigorikawa system. Because of the eight year
operation of the power plant, the fluid composi
tion has been disturbed from its initial level. Fur
ther evidence of the influence of waste water has
been obtained through tracer tests.
RELATIVE He, Ar AND N2 CONTENTS
OF GEOTHERMAL STEAM
Compared with non-condensable gas concen
trations at the initial stage of development
(Table 3), those in 1985 (Table 4) have lower
values, except for ND-5. Compositional changes
of well discharges in the first several years of
operation, coupled with the results of tracer
tests, indicate extensive mixing of waste water
with.ascending geothermal fluid. Therefore, it is
Nigorikawa geothermal system 217
Table 4. Chemical compositions of steam discharged from wells in 1985
Well Sampling datePressure of Steam fraction
in dischar e g separator (k
g /cm2G) at pressure(mole %) Total gas CO Z H2S in steamH2(mole %) N2 CH4 Ar He
ND-1
ND-3
ND-5
ND-6
ND-7
NF-1
NF-91985.10.10
1985.10.10
1985.10. 9
1985.10.11
1985.10.11
1985.10.12
1985.10.117.3
8.2
7.8
7.0
7.2
6.8
6.80.12
0.10
0.12
0.10
0.18
0.18
0.091.04
1.73
3.58
0.83
1.36
1.34
1.1397.5 0.9
97.6 0.9
98.0 1.3
97.5 0.8
97.5 1.1
97.5 1.0
97.1 0.50.0344
0.0427
0.0386
0.0878
0.0146
0.0180
0.01951.06
0.957
0.433
1.09
0.839
0.801
1.640.499
0.494
0.225
0.517
0.541
0.577
0.739x 10-4
60.9
53.8
30.2
73.8
53.7
40.3
58.1x 10-4
2.42
2.25
0.85
2.09
2.14
2.35
3.96
0. 0 1N2
Magmatic gas for
New Zealand
(Oiggenbach,1986)
Magmatic gas for
Japan(Kiyosu,1985) NF-9 \
ND-ND-1
"-1 l ND4
ND-7 " ND-5
Matsukawa
KakkondaNigorileaa
ivdcaric
front
kaida
Il 10~0k
138E 140E 142E
Air saturated water42N
40N
38N
1 0 H eCrustal
Fig. 17.A r
Relative He, Ar and N2 contents. Data of Matsukawa and Kakkonda from Yoshida (1986) .
likely that the decreasing gas concentrations are
due to the influx of degassed waste water into the
reservoir.
Although the gas concentration of hot water
is considerably affected by dilution with degass
ed waste water, the gas composition is little
affected because of the negligible gas content of
the waste water. Thus, the study of gas composi
tions of steam discharged from wells is useful for
interpretating their origin in a geothermal
system.
Figure 17 shows the relative He, Ar and N2
contents of the Nigorikawa geothermal steam.
In this figure, the gas compositions of geother
mal steam of Matsukawa and Kakkonda areas
(Yoshida, 1986) are also shown. All compositions plot along a mixing line between air
saturated groundwater and NF-9, which has the
highest contents of He and N2. Furthermore,
geothermal steam of the Matsukawa and Kak
konda systems also lie along the same mixing
line. Since the gas compositions of steam from
several geothermal systems in the region all lie
along one mixing line, there appears to be only
one high temperature endmember which mixes
with atmospheric air dissolved in groundwater.
Such mixing has been recognized in volcanic
fumaroles and geothermal wells in Japan and
New Zealand (Kiyosu, 1985; Giggenbach, 1986).
The endmember with high N2 and He contents is
derived from andesitic magmas (Giggenbach,
1986). The larger component of magmatic-deriv
218 Y. Yoshida
ed gas in the Nigorikawa system may be related
to the large scale caldera structure. Thus, deep
gas with high N2 and He contents easily ascends
along the Nigorikawa system, compared with the
other systems, consistent with a large magmatic
component to the waters (one possible interpreta
tion of the stable isotopes). Differences in the rel
ative He, Ar and N2 contents among wells of the
Nigorikawa system are due to different propor
tions of the two endmembers in the mixing proc
ess. The contribution of crustal helium or
biogenic nitrogen is negligible in this system.
SCALING PROBLEM
CaCO3 scale deposition
Plugging of production wells by CaCO3 scale
occurred soon after the power plant began
operating. The occurrence of CaCO3 scale in
each well is shown in Fig. 18. In the early stage
of the scaling problem, reaming and acid treat
ment were conducted to remove the CaCO3 scale
in the affected well. From 1984 to 1985, a scale in
hibitor (polyacrylic acid) was injected directly
below the flashing point in all production wells .
Subsequently, CaCO3 scale has been prevented,
thus solving the problem.
Figure 19 shows the relationships between
reservoir temperature and activity products ofCa" and CO3 for discharge compositions of
the initial stage (1977-1981) and during a later
stage (1985). Calculation of activities of
chemical species was carried out by the method
of Chiba (1990). Analytical data for the later
stage are listed in Tables 4 and 5.
All reservoir fluids were saturated with calcite
in 1985 (Fig. 19), such that they had the potential
to deposit CaCO3 scale during the flashing proc
ess, which causes an increase in CO3 due to the
pH increasing from C02 loss to steam. Further
evidence for CaCO3 saturation in the system
comes from the presence of calcite veins in drill
core samples; also, limestone exists in the pre
Tertiary formation. However, the fluids did not
precipitate CaCO3 because they were being dos
ed with a scale inhibitor. All well discharges
from the earlier period (Fig. 19) lie in the area of
calcite undersaturation. This apparent under
saturation with calcite is not real, but resulted
from reduced Ca concentration in the early hot
water samples following CaCO3 deposition in
the wells. Thus, caution must be used in
calculating the state of calcite saturation,
because if calcite is actively precipitating in the
well during discharges, the fluid may appear to
be undersaturated with respect to calcite (after
the fluids are recalculated to reservoir condit
ions).
Fig. 18.Total depth D-1
500
1,000
1,500
2000
2,500D-3 D-5
-16'
i I1
736MD-6 F-1 F-9
--20t-
-• 13••
.--.9 5/;.
2404ma%
2320 m
PIP CaCOs scale -133....
2,205 m ek ......
2464m
--• Point of complete circulation losses:.2d
-g 5/8
-eY2
2,340m
Configuration of CaCO3 scale of each production well (after Japan Metals and Chemicals Co ., Ltd.).
Nigorikawa geothermal system 219
N M
O
U
CZ
N No
U
•V
ftQ
Q-10,0
-12.0
-14.0C'af
°f~e
sad
GtabGO
I eo :Initial stage
• :1985
a a
a Y. a
a
s
a
158 200 258 300
Temperature ("C)
Fig. 19. Solubility of calcite and calculated activity products of geothermal water in the initial stage of the ex
ploitation, and in 1985. 0 : in initial stage; • : in 1985.
Table 5. Chemical compositions of hot waters discharged from wells in 1985
Well Sampling date Tqtz pH T-C02 Cl SO4
(°C) (mg/1) (mg/1) (mg/1). Na
(mg/1) K Ca Mg Si02 Fe Al H2S
(mg/l) (mg/1) (mg/1) (mg/1) (mg/1) (mg/1) (mg/1)
ND-1
ND-3
ND-5
ND-6
ND-7
NF-1
NF-91985.10.10
1985.10.10
1985.10. 9
1986.10.11
1985.10.11
1985.10.12
1985.10.11225 8.23 116
219 8.10 140
227 8.49 343
215 7.78 125
252 8.09 178
248 8.11 162
209 7.96 1879540 393
8590 350
4300 323
9350 369
10100 255
10100 264
9420 3895910
5380
2930
5880
6410
6360
5910815
716
416
768
923
899
73290.5
85.1
27.3
116
25.5
30.8
83.80.8
1.8
0.9
2.4
0.2
0.1
0.6486
449
500
421
704
664
3870.5
0.4
0.6
0.7
0.5
0.4
0.50.06
0.07
0.08
0.09
0.10
0.19
0.11<0.1
<0.1
0.4
<0.1
<0.1
<0.1
<0.1
Collected at atmospheric pressure
Tqtz: temperature of quartz geothermometer (adiabatic cooling)
pH: measured at room temperature
T-CO2: (H2CO3+HCO3 +C03-) as CO2
Stevensite scale deposition
Figure 20 shows the changes with time in Ca
and Mg concentrations in the ND-1 reservoir liq
uid. Mg concentration increased abruptly in
September 1988. Although scale had slowly
deposited before 1988, the precipitation rate of
scale after September 1988 accelerated in the
two-phase pipeline of ND-1 and other related
transmission lines.
The chemical composition of the scale islisted in Table -6. Scales are mainly composed of
MgO and Si02, and were identified by X-Ray
diffraction patterns to be stevensite. The
Si02/MgO ratios of scale samples from ND-1
two-phase pipeline and separators, in which fluid
temperatures are about 160°C, range from 1.92
to 2.07. Contrary to this, those from the flasher
and injection pipeline, in which fluid tempera
tures are about 120°C, have higher values of
2.69 to 2.74. This can be explained by an addi
220 Y. Yoshida
100
90
E 80 CL
70
c 0
60
ao
"' 50
a
40
30
20
10
0ND-1
o Ca
Mg
000
0 00 0
0
0O O
0
O 0
O 0
09
000 0~ 0 °
0 0
O
314111E
WX
)K )K i 31E WA M WOA% W41 )K V )E 0 I I I 10
OQD~O
O
O
* 31ME O OCO
qb 0
00
0 ,o0 O 0
O
31fO
O
*
3iE31*
31E
31E31E
31E
31E 31f
31E
31E
912 3 6
83
Fig. 20.912 3 6 9 12 3 6 9 12 3 6 9 12 3 6 9 12 3 6 912 3 6 9 12 3 6
X84 185 X86 X87 X88 89 90
Changes with time in Ca and Mg concentrations of ND-1.9 1210
9
8
7
6
5
4
3
2
1
0E
CL
0
0
U
X:
Table 6. Chemical compositions of scale samples. Sample of ND-1 well was collected in 1987,
were collected in 1989other samples
Flasher
No. 300Reinjection
line-BReinjection
line-FSeparator
No. 100Separator
No. 200ND-1 well
two-phase
Si02
Ti02
A1203
Fe203
MnO
MgO
CaO
Na20
K20
P205
H20
H20+
Si02/ MgO46.51
0.01
0.68
0.01
1.64
16.99
4.47
1.06
0.49
0.04
17.86
8.02
2.7451.09
0.01
1.21
0.12
0.91
19.03
2.20
1.25
0.51
0.03
14.71
5.35
2.6941.38
0.01
1.05
0.08
0.17
15.31
2.47
2.29
0.98
0.06
17.72
5.60
2.7050.87
0.01
1.76
0.12
1.46
25.51
1.23
1.27
0.49
0.05
11.93
5.31
1.9950.32
0.01
1.47
0.21
1.05
26.16
0.98
1.37
0.43
0.04
12.48
5.48
1.9252.14
0.01
1.34
0.28
3.18
25.14
1.44
1.19
0.51
0.03
8.91
5.85
2.07
tion of amorphous silica to stevensite at the
lower temperature.
Mg-rich clay scales have been reported for
geothermal systems of Iceland and the Philip
pines (Reyes and Cardile, 1989; Gunnlaugsson
and Einarsson, 1988; Kristmannsdottir et al.,
1988). Stevensite scale deposition in the
Nigorikawa geothermal system relates to an increase of Mg concentration in well discharges.
The increasing Mg concentration also correlates
with a decrease in enthalpy of the reservoir fluid.
An abrupt increase in Mg concentration occur
red in September 1988 (Fig. 20). During this
period, drilling was taking place near ND-1, and
a significant amount of drilling fluid was lost
into the formation. Although drilling was com
Nigorikawa geothermal system 221
pleted, stevensite is still precipitating. Therefore,
it is likely that a low enthalpy fluid entered the
ND-1 reservoir as a result of a changing flow pat
tern of the geothermal fluid. Mg in the low en
thalpy fluid was contributed from shallow ther
mal water enriched in Ca and Mg, similar to the
Nigorikawa hot spring waters.
CONCLUSIONS
Thermal waters of hot springs in and around
the Nigorikawa Basin and waters discharged
from wells may be divided into four groups on
the basis of their relative CI and SO4 contents
(hot springs outside the Basin, hot springs inside
the Basin, hot water derived from faults in the
pre-Tertiary Kamiiso Group, and hot water
derived from the fractured zone of the caldera
wall). Although S04 in the thermal water prob
ably comes from dissolution of anhydrite at
depth in host-rock formations, there are no
chemical trends indicating this dissolution. The
Nigorikawa hot spring water is formed by mix
ing of hot water ascending along caldera wall
fractures and Ca-rich zero chloride water; the lat
ter has a steam-heated origin. In the mixing proc
ess, the SO4 concentration decreases by dilution
and by shallow anhydrite deposition.
The relationship between the Cl concentra
tion and enthalpy of the initial stage reservoir
fluid indicates dilution by a steam-heated water
with temperature ranging from 150 to 200°C.
The variable temperature of the diluent relates to
a variable degree of steam-heating, which ac
counts for the distinctive Ca and SO4 concentra
tions between wells, and is due to a variation in
permeabilities. In spite of local differences in
chemical composition, isotopic data indicate
that the geothermal fluid is formed by simple
mixing of heated meteoric water with deep hot
water, with the latter having a large component
of magmatic and/or altered sea water.
Relative He, Ar and N2 content of the
Nigorikawa geothermal fluids indicate that these
gases are simple mixtures of magmatic-derived
gas and atmospheric air dissolved in ground
water. Since the Mori power station began opera
tion, the chemical compositions of well
discharges have changed. Based on trends in the
Cl-enthalpy relationship, three endmembers, i.e.
original deep fluid, zero chloride shallow steam
heated water and reinjected waste water, have
been identified in the geothermal system. Reser
voir fluids of each well are mixtures of these end
members, and recently, the contribution of the
zero chloride shallow steam-heated water is in
creasing.
The reservoir liquid of the Nigorikawa
geothermal system is saturated with calcite,
resulting in CaCO3 scale deposition in wells as a
result of flashing. A prevention system using a
CaCO3 scale inhibitor was completed in 1985,
solving the CaCO3 scaling problem. On the other
hand, stevensite scale has recently precipitated in
the two-phase and hot water pipelines related to
well ND-1. This well was affected by an incur
sion of Mg-rich low enthalpy water.
Acknowledgments-I thank the members of the Ex
ploration Division of JMC Geothermal Research and
Development Co., Ltd., who collected much of the
data presented here. The information on distribution
of anhydrite in wells was contributed by Mr. R.
Komatsu. I also thank Japan Metals and Chemicals
Co., Ltd. and Dolman Geothermal Energy Co., Ltd.
for their help in sample collection and for supplying
much of the data. Critical comments by Dr. J. W.
Hedenquist of the Geological Survey of Japan helped
to improve an earlier version, of the manuscript.
REFERENCES
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Fukutomi, T., Fujiki, T., Sugawa. A., Ohtani, K.,
Wada, A. and Tokunaga, E. (1963) On the hot spr
ings of Nigorikawa in southern Hokkaido.
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Giggenbach, W. F. (1986) The use of gas chemistry in
delineating the origin of fluids discharged over the
Taupo Volcanic Zone. Proceedings of Symposium
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Gunnlaugsson, E. and Einarsson A. (1988)
Magnetism-silicate scaling in mixture of geothermal
222 Y. Yoshida
water and deaerated fresh water in a district heating
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Ide, T. (1982) Geology in the Nigorikawa geothermal
field, Mori-machi, Hokkaido, Japan. Geotherm.
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Ide, T. and Doi, N. (1982) Neogene formations
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Igarashi, T., Sato K., Ide, T., Nishimura, S. and
Sumi, K. (1978) Geological investigation of
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Kristmannsdottir, H., Olafsson, M. and
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Matsubaya, 0., Sakai, H., Ueda, A., Tsutsumi, M.,
Kusakabe, M. and Sasaki, A. (1978) Stable isotope
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steam from Showashinzan Volcano, Hokkaido,
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Chemical and isotopic compositions of fumarolic
gases from Kuji-Iwoyama, Kyushu, Japan.
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Nagao, K., Takaoka, N. and Matsubayashi, O. (1979)
Isotopic anomalies of rare gases in the Nigorikawageothermal area, Hokkaido, Japan. Earth Planet.
Sci. Lett. 44, 82-90.
Reyes, A. G. and Cardile, C. M. (1989) Characteriza
tion of clay scales forming in Philippine geothermal
wells. Geothermics 18, 429-446.
Sato, H. (1969) Newly obtained 14C data concerning
volcanic activity in Hokkaido. Chishitsu News 178,
30-35 (in Japanese).
Urakami, K. and Nishida, Y. (1977) Heat discharge
measurement and geophysical prospecting at
Nigorikawa Basin, northern part of Komagatake.
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study on hot springs. -Report No. 44, Nigorikawa
hot spring. Nihon Kagaku Zasshi 80, 866-871 (in
Japanese).
Yoshida, Y. (1981) Chemical study on deep hot water
of the Nigorikawa geothermal field. Geotherm.
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Yoshida, Y. (1982) Study on the hydrothermal system
in the Nigorikawa Basin, Hokkaido. Onsenkagaku
33, 24-36 (in Japanese).
Yoshida, Y. (1983) Calcium carbonate saturation of
hot water in the Nigorikawa geothermal field. Ex
tended Abstracts of the 4th Internaltional Sym
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545-548.
Yoshida, Y. (1986) He, Ar and N2 concentrations in
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651.
Yoshida, Y. (1990) Chemical study on the hot springs
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|
Yoshida (1991) geochemistry of nogorikawa geothermal system.txt
|
Gondwana Research, V 6, No. 4, pp. 687-698.
0 2003 International Association for Gondwuna Research, Japan
ISSN: 1342-937X
Jurassic Dextral and Cretaceous Sinistral Movements Along the
Hida Marginal Belt
Kazuhiro Tsukada
The Nagoya University MUS~ZLWI, Nagoya 464-8601, Japan
(Manuscript received June 28, 2002; accepted March 25, 2003)
Abstract
Thc Hida marginal bclt (HMB), which consists of various kinds of fault-bound blocks, is located bctwccn the continental
inassif of the Hida bclt and the Mesozoic accrctionary cornplcx of the Mino belt in Ccntral Japan. Detailcd field
investigation reveals that thc HMB had grown through the two diffcrcnt movements, i.c., Jurassic dextral and Crctaceous
sinistral movcmcnts. The Jurassic dextral ductile shcar zones run in the southcrn marginal part of thc Hida bclt and the
northern part of thc HMB, whereas the Cretaceous sinistral cataclastic shear zones occur in the southcrn part of the
HMB and the northern marginal part of the Mino belt. Geologic map and field cvidcncc sccin to suggest that the
Jurassic dcxtral movcmcnt form thc fault-bound blocks of the HMB to form the basic structure of the Hida marginal
bclt, i.c., formation of thc ‘proto-HMB.’ Following thc dextral movcment, thc sinistral onc rcstructurcd the ‘proto-HMB’
to complctc thc prcscnt feature of thc Hida marginal belt. The Cretaceous sinistral movcment might result in the
sinistral collision betwccn thc proto-HMB and the Mino belt.
Key words: Hida marginal belt, shcar zoncs, Jurassic and Cretaceous movements, southwest Japan.
Introduction
The geology of the East Asian contincntal margin
developed through various processes such as accretion,
collision and strilte-slip tectonic movement. Accretion of
trench-fill sediments and oceanic plate cover, as well as
collision of terraiies with strike-slip tectonic movement
created the basic geotectonic framework of SW Japan.
Features of the accretionary complexes of SW Japan have
been studied and various tectonic models for their
formation have been presented (e.g., Wakita, 2000).
However, the process of the collision and the role of the
strike-slip tectonic movement in the evolution of SW Japan
have not been precisely explained.
In SW Japan, the Hida, Sangun, Akiyoshi, Maizuru,
Ultra-Tamba and Mino bclts are distributed from north to
south in the Inner Zone of SW Japan, i.e., on the north of
MTL (Fig. 1). The Sangun, Akiyoshi, Maizuru, and Ultra-
Tamba belts are widely exposed in the western part of
SW Japan, but are absent in the eastern part of SW Japan.
The Hida marginal belt, which includes fault-bound blocks
derived from the Sangun, Akiyoshi and Maizuru belts and
Paleozoic shelf facies rocks (e.g., Chihara et al., 1979) is
distributed in a narrow zone between the Hida and Mino
belts in eastern part of SW Japan (Fig. 1). The blocks
derived from the Paleozoic shelf facies rocks can be divided into the two types based on their lithostratigraphy (Waltita
et al., 2001). Judging from the distribution of the belts of
the Inner Zone, the Hida marginal belt may be interpreted
as a tectonic zone formed by the collision of the Hida and
Mino belts. In order to examine this hypothesis, analysis
of rock distribution and fabrics of shear zones in and
around the Hida marginal belt is necessary. In this paper,
the stratigraphy, the rock distribution and the fabrics in
the shear zones in the Fukuji-Takayama area, including
the Hida belt, the Hida marginal belt and the Mino belt
central Japan, are discussed in order to consider the
evolution of thc Hida marginal belt.
Geological Setting of Southwest Japan
The Japanese Islands are geographically divided into
SW Japan and NE Japan by Itoigawa-Shizuoka Tectonic
Line, and SW Japan is subdivided into the Inner Zone
and the Outer Zone by the Median Tectonic Line (Fig. 1).
The Inner Zone of SW Japan is composed of the following
belts, Hida belt, Hida marginal belt, Mino belt, Ryoke belt,
Altiyoshi belt, Sangun belt, Maizuru belt, Ultra-Tamba belt
and Nagato tectonic belt (Wakita, 1989; Wakita et al.,
1992, Fig. 1). The Hida and Sangun belts consist mainly
of Paleozoic and Mesozoic metamorphic rocks, and the
Ryolte belt is of the metamorphosed accretionary complex Gondw ana
, Research GR
688 K. TSUKADA
of the Mino belt (e.g., Geological Survey of Japan, 1992).
The Akiyoshi and Ultra-Tamba belts are characterized by
Permian accretionary complexes, whereas the Mino belt
is Jurassic (e.g., Geological Survey of Japan, 1992). The
Maizuru belt is an upper Paleozoic island arc system
covered by Upper Paleozoic to Lower Mesozoic strata
(e.g., Geological Survey of Japan, 1992). The Nagato
tectonic belt is considered to be the western extension of
the Hida marginal belt. Most of these belts occur as
subhorizontal or gently-northward-dipping thin tectonic
units and form a huge pile of nappes in the western part
of the Inner Zone of SW Japan (e.g., Isozaki et al., 1990).
The Sangun, Akiyoshi, Maizuru, and Ultra-Tamba belts
are widely exposed in the western part of SW Japan,
whereas these belts are absent and the Hida marginal belt
is relatively narrowly distributed between the Hida and
Mino belts in the eastern part of SW Japan (Fig. 1).
It is generally considered that the Jurassic accretion of
the Mino belt formed the nappe system in the western
part of the Inner Zone of SW Japan, and stacking of nappes
of the Sangun, Akiyoshi, Maizuru and Ultra-Tamba belts
was enhanced during Jurassic time (e.g., Hayasaka, 1990;
Ichikawa, 1990). This means that the Sangun, Akiyoshi,
Maizuru, Ultra-Tamba and Mino belts must have been amalgamated by the Jurassic. In this paper, the term ‘Outer
belt’ will be used for the amalgamated Sangun-Akiyoshi-
Maizuru-Ultra-Tamba-Mino belt for convenience.
Paleomagnetic reconstruction indicates that the
Japanese Islands were arranged in a northeast direction
before the opening of the Sea of Japan in the Middle
Miocene (Otofuji et al., 1985, Fig. 1). Most of the belts in
Japan did not extend to Korea, but run on the east of the
Korean Peninsula trending north to northeast, although
part of the Hida belt may have extended to the Gyeonggi
or Yeongnam Massif of Korea (Otoh et al., 1999). This
means that the Hida belt was likely a part of the China
block before the opening of Sea of Japan, but the rocks of
the other belts in Japan have different origins and
evolutionary histories (Otoh et al., 1999).
Geological Framework of the Hida Marginal
Belt
The Hida marginal belt was first defined by Kamei
(1955 in Nozawa, 1978) as a complex zone dividing the
continental massif of the Hida belt and the ‘Paleozoic
geosynclinal facies’ rocks of the Mino belt (Fig. 2). It has
been dealt with as follows: a late Paleozoic geosyncline
Legend
ETI] I2 Hidd belt
@# Sdngun belt
Akiyo5hi belt
Maizuru belt
Ultra-Tarnba belt
Mino belt
(Ryoke belt included)
Hidd marginal belt and
Nagato tectonic belt FTj Rocks of the Outer Zone
Fig. 1. Tectonic map of southwest Japan before the opening of the Sea of Japan. The tectonic subdivision of Wakita et al. (2001) and Otoh et al.
(1999) is followed. Present features of the Japanese Islands are shown with pale lines. MTL, ISTL, TTL, HMB and NTB show the Median
tectonic line, Itoigawa-Shizuoka tectonic line, Tanakura tectonic line, Hida marginal belt and Nagato tectonic belt, respectively.
Gondwana Research, V. 6, No. 4, 2003
DEXTRAI, AND SINISTRAL MOVEMENTS HIDA MARGINAL BELT 689
which evolved into a narrow tectonic zone in late Jurassic
time (Kamei, 1955 in Nozawa, 1978), a serpentinite
mklange zone caused by the tectonics forming the ‘Hida
Nappe’ (Komatsu, 1990; Chihara and Komatsu, 1982;
Soma and Kunugiza, 1993) that overthrust from the Hida
belt onto the ‘Outer belt,’ and a complex zone formed by
several stages of movements from Permian to Jurassic
times (Kimura et al., 1993).
The Hida marginal belt consists of various kinds of fault-
bound blocks; e.g., blocks derived from the ‘Outer belt’
such as the Sangun, Altiyoshi and Maizuru belts
(e.g., Chihara et al., ‘1979; Komatsu, 1990; Nishimura,
1990; Takeuchi, 1998; Kawai and Takeuchi, 2001), and
blocks derived from the Paleozoic shelf-facies rocks
(Fig. 3). The blocks derived from the Paleozoic shelf-facies
rocks can be divided into the following two types based
on their lithostratigraphy: (1) blocks composed mainly of
Devonian limestone, Carboniferous limestone and Permian
clastic and pyroclastic rocks (named as the Fukuji-type
block), (2) blocks composed mainly of Devonian felsic
tuffaceous clastic rocks, Carboniferous mafic pyroclastic
rocks and Permian clastic rocks (named as the Moribu- type block) (Fig. 3). In the eastern part of the Hida
marginal belt, equivalents of the Sangun, Akiyoshi and
Maizuru belts, which trend north to northeast and dip
west, have been overthrust by the Hida belt with shear
plane trending northeast and dipping west (e.g., Komatsu,
1990). In other places, the rocks of the Hida marginal
belt trend east and dip subvertically. The contact between
the Hida marginal belt and the Mino belt is a shear zone
with foliated fault rocks. Foliation in the shear zone
generally trends east and dips southward to subvertically.
In this paper, the focus is on the shear zones in and
around the Hida marginal belt and the blocks derived from
the Paleozoic shelf-facies rocks. The geological setting of
the Fukuji-Takayama area, which is the type locality of
the Fukuji- and Moribu-type blocks is described.
Geological Description of the Fukuji-
Takayama Area
Structure, stratigraphy and paleobiogeography
In the Fukuji-Takayama area, the Moribu-type block,
the Lower Cretaceous beds of the Tetori Group (Maeda,
Mcsozoic to Cenomic
Funatsu Granite Rock\
of the Htda
marginal belt
Schist 0 Gneijs / Fault a cover rock4
1 km
Fig. 2. The distribution of the gneiss, schist, and Funatsu Granite in the Hida belt and rocks of the Hida marginal belt (modified from Nozawa, 1977
and Hiroi, 1981). ISTL-Itoigawa-Shizuoka tectonic line. See figure 1 for the locality.
Gondzuana Rcwauch, K 6, No. 4, 203 Sea of Japan
690 K. TSUKADA
Hida belt Hida marginal belt Mino belt
N -
Tertiary
Moribu- Akiyoshi belt
equivalent of
the Saiigun
belt Origindl ruck\ ot
I'recarnhrian quivalent of'
he Maizuru
Nelt
Lzl Fukuji-
type
-
Legend
: ; Felsic volcanic rocks and
L L L tuffxeous clastic rocks
.,.',.,.'...'.... Post-Triassic shallow to
non-marine sediments ,',:,'.:,..'..
........... ...... ... ..........
....
Granite
$-@$ Siliceous shale
:{${$$: Conglomerate
Sandstone
and mudstone I,........
........ 1111111 Chert
..l^^ ,.**.
..A*b~A*. Basalt
,^__
Gabbro Serpentinite
Limestone
Tuffaceous S rnklange
5% clastic rocks
? Ageunknown - Unconformity
Fig. 3. Stratigraphic summary of the Hida and Mino belts and the Hida marginal belt. The Mino belt is represented by the reconstructed ocean
plate stratigraphy of the Funafuseyama unit (Wakita, 2000). Although radiometric ages of the Funatsu Granite show one peak at around
180 Ma, minor plutons have the radiometric ages of 200-250 Ma (Shibata and Nozawa, 1982; Ota and Itaya, 1989). G-Group, Gr-Granite,
E-Early, M-Middle, L-Late, P-Paleogene, N-Neogene.
1958) and the Fukuji-type block are distributed in the Hida
marginal belt from north to south between the Hida and
Mino belts (Fig. 4). Although the pre-Cretaceous rocks in
the area generally trend east to northeast and steeply dip
northward or subvertically, they trend north and steeply
dip westward around the Moribu (Fig. 4). The Moribu-
type block, the Lower Cretaceous beds of the Tetori Group
and the Fukuji-type block are bounded by shear zones
with each other. Some shear zones can be seen in the
Moribu-type block and the northern marginal part of the
Mino belt (e.g., Sasaki et al., 2001; Fig. 4). The Moribu-
type block is subdivided into several fault-bound blocks
by the shear zones. A shear zone bound the Moribu-type
block and the Upper Triassic Tandodani Formation around
the Hongo (Tsukada et al., 1997). The Moribu-type block in the study area is composed
of the Devonian Rosse (mainly felsic tuff), Carboniferous
Arakigatva (mainly mafic volcanic rocks), and Lower
Permian Moribu (mainly clastic rocks) formations
(e.g., Isomi and Nozawa, 1957; Igo, 1990; Tazawa et al.,
2000; Wakita et al., 2001, Fig. 5). The Arakigawa
Formation conformably or unconformably underlies the
Moribu Formation (Horikoshi et al., 1987; Wakita et al.,
2001). The Fukuji-type block is composed of the Upper
Silurian (?> to Devonian Yoshiki (felsic tuffaceous clastic
rocks), Devonian Fukuji (mainly limestone), Carboniferous
Ichinotani (mainly limestone), Lower Permian
Mizuyagadani (mainly clastic rocks), and Lower Permian
Sorayama (mainly mafic volcanic rocks) formations
(e.g., Igo, 1990; Tsukada and Takahashi, 2000; Wakita
Gondwana Research, V 6, No. 4, 2003
DEXTRAL AND SINISTRAL MOVEMENTS HIDA MARGINAL BELT 691
et al., 2001, Fig. 5). Although most of the formations in
the Fukuji-type block are presently in fault contact with
each other, they are interpreted to form a primarily
conformable or unconformable succession, because the
Permian formations include abundant clasts derived from
the Devonian to Carboniferous formations (Igo, 1990;
Tsukada and Takahashi, 2000; Wakita et al., 2001).
Sandstone and mafic tuff of the Moribu-type block as well
as gneiss of the Hida belt are intruded by the Funatsu
Granite (Figs. 2, 3, 4). K-Ar and Rb-Sr radiometric ages
of the Funatsu Granite show one peak at around 180 Ma,
and minor plutons shdw radiometric ages of 200-250 Ma
(Ota and Itaya, 1989; Shibata and Nozawa, 1982). The
Middle Jurassic to Lower Cretaceous Tetori Group
unconformably overlies the Funatsu Granite (Figs. 3, 4).
The Mino belt in the area is composed largely of
melange including various kinds of slabs and blocks
(Wakita, 1988; Waltita, 2000) (i.e., sandstone, mudstone,
felsic tuff, bedded chert, limestone, and mafic volcanic
rocks) set in a muddy matrix (Fig. 2). Some slabs in the
melange are enormous, and the largest one reaches 1 km
in width and extends as long as 5 km. Limestone blocks
yield Carboniferous to Permian fusulinaceans (Isomi and
Nozawa, 1957; Kojima, 1984). Permian to lower Middle
Jurassic radiolarians are obtained from chert blocks (Kojima, 1984). The muddy matrix and felsic tuff blocks
yield Middle Jurassic radiolarians (Kojima, 1984).
Judging from the fossil data, the age of accretion is
considered to be Middle Jurassic. The all pre-Lower
Cretaceous rocks of the area are unconformably covered
by undeformed uppermost Cretaceous to Quaternary
volcanic rocks (Fig. 3). In the eastern part of the area, the
rocks of the Hida marginal belt and the Mino belt
underwent contact metamorphism by the intrusion of
earliest Tertiary to Quaternary granitoids (Fig. 4).
The Lower Permian beds of the Moribu-type block yield
brachiopods forming a mixed fauna of Boreal and Tethyan
species which have close faunal affinity to those from Inner
Mongolia (eg, Tazawa, 1991, Fig. 5). The fusulinacean
genus Monodiexodina, which commonly occurs on
northeast China, Mongolia, and Siberia is also found from
the Lower Permian beds of the Moribu-type block (e.g.,
Ishii et al., 1985; Tazawa et al., 1993, Fig. 5). Coeval beds
of the Fukuji-type block, in contrast, yield a typical Tethyan
fusulinacean fauna having close faunal affinity to that from
the South China (Tsukada et al., 1999; Tsukada and
Takahashi, 2000, Fig. 5). Lower Cretaceous beds of the
Tetori Group in the area yield a ‘Tetori-type’ flora having
a close affinity to that of North China and Siberia (Kimura,
1987; Maeda, 1958).
Hida I Cwer rocks Funatsu Granite Pukuji-type
Tcrtiary to Quarternary granites Accretionary complex Moribu-type marginal belt
@ Rocks of the
other blocks [3 Tetori Group
%%%%W Dextral shear zones ~~~~~ Sinistral shear zones ~ Fault
Fig. 4. Simplified geological map of the Fukuji-Takayama area showing the distribution of shear zones. See figure 2 for the locality. HMB-Hida
marginal belt.
Gowdzuui~a Keseurch, V 6, No. 4, 2003
692 K. TSUKADA
Fukuji-type Moribu-type
*------- ............... Mixed fauna of Tethyan and Boreal
I unconformitv ? 1 unconformitv ? I
***n***fi*** ........... ...........
.rakigawa F.: ........... ........... ........... ........... ........... * * * * .-. .- * -7 * *A n n ............. ............. ............. ............. ............. ............. ............. ............. .............
I I
Legend
Mafic to intermediate pyroclastic rocks
and tuffaceous clastic rocks
Felsic tuff and tuffaceous clastic rocks Conglomerate
Limestone Fig. 5. Schematic columnar sections of
the Fukuii-type and Moribu-
Sandstone and mudstone
Description of shear zones
Ductile shear zones occur in the southern margin of
the Hida belt and the Moribu-type block (Fig. 4). The
ductile shear zones in the Hida belt cut the Funatsu Granite
to form augen gneiss (Fig. 6a). The shear zones in the
Funatsu Granite are composed of mylonite, and have a
northeast-trending subvertical foliation with a
subhorizontal lineation. In the mylonite, fragments of
quartz are strongly deformed and dynamically recrystallized. type blocks -in the Fukuji-
Takayama area. F-Formation.
The foliation in the ductile shear zones is mainly defined
by dimensional preferred orientation of micas. In meso-
to microscopic observation, various kinds of shear sense
indicators showing a dextral sense of shear, e.g., shear
band (White, 1979), S-C fabric (Berthe et. al., 1979) and
asymmetric pressure shadows, are observed in the mylonite.
The shear zones in the Moribu-type block are composed
mainly of schistose rocks derived from mafic tuff and
clastic rocks, and have bedding-parallel foliation and
Gondwana Research, V 6, No. 4, 2003
DEXTRAL AND SINISTRAL MOVEMENTS HIDA MARGINAL BELT 693
subhorizontal to subvertical lineation. The shear zones
trend east and steeply dip northward or subvertically
around the Kamihirose and Hongo, whereas, they trend
north and steeply dip westward around the Moribu
(Fig. 4). Toward the east, around the Moribu, the schistose
rocks gradually shift into the foliated cataclasite. In the western part of this area, the schistose rocks are
dynamically recrystallized and fragments of quartz show
wavy extinction (Fig. 6b). The foliation in the Moribu-
type block is defined by concentration and dimensional
preferred orientation of metamorphic minerals such as
chlorite and muscovite (Fig. 6b, c). In meso- to microscopic
Fig. 6. Photographs showing the meso- and microstructure of the fault rocks of the dextral shear zones in the Fukuji-Takayama area. All
photomicrographs show the sections cut perpendicularly to the foliation and subparallel to the lineation. (a) Outcrop of the mylonite
developed in the Funatsu Granite. The mylonite is composed largely of porphyroclasts of K-feldspar surrounded by fine-grained quartz and
plagioclase. (b-d) Photomicrographs of the sheared rocks in the Moribu-type block. (b) Schistose rock derived from sandy mudstone of the
Moribu Formation. Quartz grains are dynamically recrystallized and show wavy extinction. Concentration and dimensional preferred
orientation of muscovite (upper white part) define the foliation. (Crossed polars) (c) Schistose rock derived from mudstone of the Moribu
Formation. Asymmetric pressure shadows on intensely elongated quartz grains (white grains) show a dextral sense of shear. Muscovite
(dark part) is arranged parallel to the foliation. (Plane polarized light) (d) Schistose rock derived from mafic tuff of the Arakigawa
Formation. Asymmetric structures clearly indicate a dextral sense of shear. (Plane polarized light) (e) Metamorphosed sheared rock. The
rock is intruded by the Funatsu Granite, and is metamorphosed at least into the epidote-amphibolite facies. The rock is composed largely
of actinolite, hornblende, epidote and plagioclase. Asymmetric tail on a grain in the middle part (black arrow) shows a dextral sense of
shear (Plane polarized light).
Gondwana Research, K 6, No. 4, 2003
694 K. TSUKADA
observation, various kinds of shear sense indicators
showing a dextral or top-to-the-east sense of shear, e.g.,
asymmetric pressure shadows and asymmetric folds, are
observed in the sheared rocks (Fig. 6c, d). The sheared
rocks in the Moribu-type block are intruded by the Funatsu
Granite and are metamorphosed at least into the epidote- amphibolite facies in the eastern part of this area, around
Fukuji (Fig. 6e).
There is a remarkable shear zone between the Fukuji-
type block and the Lower Cretaceous clastic rocks of the
Tetori Group (Fig. 4). The shear zone is commonly several
meters wide. The rocks in the shear zone are composed
Fig. 7. Photographs showing the meso- and microstructure of the fault rocks of the sinistral shear zones in the Fukuji-Takayama area. All photographs
show the sections cut perpendicularly to the foliation and subparallel to the lineation. (a-d) The sinistral shear zone along the boundary
between the Sorayama Formation (Fukuji-type block of the HMB) and the Tetori Group. (a) Polished surface of cataclasite derived from
mudstone of Tetori Group. Abundant elongated sandstone fragments (white arrows) are contained in a matrix of foliated mudstone.
(b) Polished surface of foliated cataclasite derived from sandstone of the Tetori Group. The rock has the P-Y fabric and the o-type asymmetric
structure (‘0’ in the figure) indicating a sinistral sense of shear. The P-surface has a spacing interval of several 100 pni and changes into the
Y-surface at its termination. ‘Y’ and ‘P‘ in the figure show the directions of Y- and P-surfaces, respectively. (c) Photomicrograph of foliated
cataclasite derived from sandstone of the Tetori Group showing the R1-type Riedel shear plane (black arrow). (d) Photomicrograph of
cataclasite derived from sandstone of the Tetori Group showing drag of a broken grain (small arrow). Large arrow shows a shear plane
parallel to the foliation. (Crossed polars) (e) Polished surface of foliated cataclasite derived from clast-bearing mudstone of the Mino belt
including limestone clasts. Limestone clasts (white part) are intensely elongated. The R1-type Riedel shear plane (black arrows) and
asymmetric folds can be seen in the rock.
Gondwana Reseavch, V 6, No. 4, 2003
DEXTRAL. AND SINISTRAL MOVEMENTS HIDA MARGINAL BELT 695
Fig. 8. Distribution of sinistral she3r zones in the northern marginal part of the Mino belt. The stereograms show the strike and dip of foliations and
the trend of lineations. See figure 4 for the locality. HMB-Hida marginal belt, R-River.
mainly of fissile foliated cataclasite derived from sandstone
and mudstone of the Tetori Group (Fig. 7a). Rocks in the
shear zone are not recrystallized and they have a
northeast-trending subvertical foliation with a
subhorizontal lineation. The foliation is subparallel to the
boundary between the Tetori Group and the Fukuji-type
block. The cataclasite commonly contains elongated
irregular- or columnar-shaped sandstone fragments in a
matrix of foliated mudstone (Fig. 7a). The foliation in the
muddy matrix is defined by dimensional preferred
orientation of fine-grained micas and clay minerals.
Foliated cataclasite derived from the sandstone of the
Tetori Group is also observed in the shear zone. The
foliation in the sandstone cataclasite is defined by thin
layers of fine-grained micas, clay minerals and opaque
minerals. Paralell foliation can be seen in the sandstone
cataclasite with a spacing of several tens of microns to
several millimeters (Fig. 7b). The fragments in the
sandstone cataclasite which have parallel foliation are
intensely elongated. In meso- to microscopic observation, various kinds of shear sense indicators showing a sinistral
sense of shear, e.g., Riedel shear of R1-type (Logan et al.,
1979), P-Y fabric (Rutter et. al., 1986), asymmetric
pressure shadows and asymmetric folds, are commonly
observed in the foliated cataclasite (Fig. 7b, c, d).
Some shear zones run in the northern marginal part
of the Mino belt (Sasaki et al., 2001; Niwa et al., 2002,
Figs. 7e, 8). Shear zones are commonly several meters to
500 m wide. They are east-trending with a vertical to
south-dipping foliation and subhorizontal or south- to
southwest-plunging lineations in the western part of the
area (Fig. 8). The foliation in the eastern part of the area
trends northeast and dips northwest, and the lineation
plunges west to southwest (Fig. 8). The trends of the shear
zones are parallel to the boundary between the Hida
marginal belt and the Mino belt (Fig. 4). The shear zones
are composed mainly of foliated cataclasite derived from
melange of the Mino belt. The cataclasite commonly
includes lenticular fragments of chert, sandstone,
limestone and mafic volcanic rocks in foliated muddy
Gondwnna Research, V. 6, No. 4, 2003 Upper Cretaceous to Quaternary
Rocks of the Hida marginal belt
Rocks of the Mino Belt
Shear zone
Foliation
Lineation
696 K. TSUKADA
matrix. The foliation in the muddy matrix is defined
by thin dark layers of fine-grained micas and clay minerals.
In some places, parallel foIiation can be observed in the
muddy matrix with a spacing of several microns to several
tens of microns (Fig. 7e). The clasts in the muddy matrix
exhibiting parallel foliation are intensely elongated in
general. Most of the cataclasites are not recrystallized,
but rare ductile shear zones (about a few hundred microns
in wide) are found at the northern edge of the Mino belt.
In the ductile shear zones, fragments of quartz are strongly
deformed and dynamically recrystallized. R1-type Riedel
shears (Logan et al., 1979) and P-Y fabrics (Rutter et. al.,
1986) clearly showing sinistral sense of shear are seen in
the cataclasite (Fig. 7e). Under the microscope, various
kinds of shear sense indicators showing a sinistral or
top-to-the-east sense of shear, e.g., asymmetric folds and
asymmetric pressure shadows, are commonly observed
in the cataclasite. Sasaki et al. (2001) mentioned that
the cataclastic shear zones in the north of Hiyomo
(Fig. 8) together form a sinistral strike-slip imbricate
fan.
Discussion and Conclusions
Timing of the shearing
In the Fukuji-Takayama area, the dextral and sinistral
shear zones are recognized. The dextral shear zones in
the Hida belt cut the Funatsu Granite to form augen gneiss.
The dextral shear zones in the Moribu-type block are
intruded by the Funatsu Granite, and limit the distribution
of the Upper Triassic Tandodani Formation (Fig. 4).
Magnetic foliation and lineation subparallel to the foliation
and lineation in the dextral shear zones in the rocks of
the Hida belt and the Moribu-type block occur in the
undeformed Funatsu Granite (Otoh et al., 1996). These
facts suggest that the dextral shearing took place toward
the Jurassic.
A sinistral shear zone occurs along the boundary
between the Tetori Group and the Fukuji-type block in
the Fukuji-Takayama area. Maeda (1958, 1959) reported
some bivalves and ‘Tetori-type’ plant fossils from the
Tetori Group of the area, and correlated the fossil-bearing
beds to the Lower Cretaceous beds of the Tetori Group of
the type locality. The sheared rocks in the Fukuji-
Takayama area are unconformably covered by undeformed
volcanic rocks. The lower part of the volcanic rock
succession yields the Maastrichtian palynoniorphs
(Kasahara, 1979). The sheared rocks are intruded by
undeformed granitoids dated around 64 Ma (Harayama,
1990). Hence, it is obvious that the sinistral shearing lasted
after the early Cretaceous and had finished by the latest
Cretaceous. Development of the Hida Marginal Belt
The lithostratigraphical and paleobiogeographical data
strongly suggest that the Moribu-type and Fukuji-type
blocks were formed in different regions at least by the
Middle Permian. The Funatsu Granite intrudes into the
Moribu-type block as well as the gneiss of the Hida belt
(Figs. 2, 3, 4). The Middle Jurassic to Lower Cretaceous
Tetori Group unconformably overlies the Funatsu Granite
(Figs. 3,4). This indicates that the Moribu-type block must
have been juxtaposed against the Hida belt by the Jurassic.
The timing of juxtaposition of the Fukuji-type block against
the Hida belt and the Moribu-type block is unknown
because there is no data suggesting the Mesozoic position
of the Fukuji-type block. Maeda (1961) mentioned that
the Fujikuradani Formation, which might be correlated
with the Ichinotani Formation of the Fukuji-type block is
unconformably overlain by the Middle Jurassic beds of
the Tetori Group in the western part of the Hida marginal
belt. Given that this view is correct, the Fukuji-type block
may have been juxtaposed against the Hida belt by the
Middle Jurassic.
The Jurassic dextral shear zones run in the northern
part of this area, i.e., Hida belt and Moribu-type block,
whereas the Cretaceous sinistral shear zones occur in the
southern part of this area and bound the Tetori Group,
the Fukuji-type block and the Mino belt (Fig. 4). Otoh et
al. (1999) mentioned that the dextral shear zones in this
area resulted from the northward drifting of the
continental side. Geological map and field evidence seem
to suggest that the Jurassic dextral movement caused the
fragmentation of the Moribu-type block to form the basic
structure of the Hida marginal belt, i.e., formation of the
fault-bound blocks.
Following the dextral movement, the sinistral
movement took place along the eastern margin of Asia in
the Cretaceous (e.g., Ozawa 1987; Tashiro, 1994; Otoh
and Yanai, 1996). The sinistral shearing in this area was
presumably attributed to this Cre t ace ous sinj s tral
movement. It is likely that the sinistral shearing
restructured the ‘proto-Hida marginal belt’ to complete
the present feature of the Hida marginal belt. The
Cretaceous sinistral movement might result in the sinistral
collision between the proto-Hida marginal belt and the
Mino belt.
Acknowledgments
I wish to thank Prof. M. Adachi, Associate Prof. M.
Takeuchi, and Associate Prof. H. Yoshida of Nagoya
University for helpful discussion and advice. I am indebted
to Prof. S. Kojima of Gifu University, Dr. K. Wakita of
Geological Survey of Japan, Associate Prof. S. Yamakita
Goiidwann Research, I/: 6, No. 4, 2003
DEXTRAL AND SINISTRAL MOVEMENTS HIDA MARGINAL BELT 697
of Miyazalti University, Prof. Emeritus H. Igo of University
of Tsultuba and participant of the field worltshop
(FW-C1) of ISRGA of 2001 for valuable discussion. I would
liltc to thank Mr. M. Niwa, Ms. K. Hotta and Mr. S. Tanaka
of Nagoya University for valuable suggestion during the
course of this work. The work was supported by Grant-
in-Aid of Fukada Geological Institute and Grant-in-Aid of
Fundamental Scientific Research (Nos. 10003433 and
11 740278) from the Ministry of Education, Science and
Culture, Japan. Special thanks go to Prof. J. Aitchison of
University of FIong Kong and Associate Prof. S. Otoh of
Toyama University fo? critical reading- of the manuscript.
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Available online at www.sciencedirect.com TECTONOPHYSICS SCIENC DIRECT ELSEVIER Tectonophysics 403 (2005) 59-75 www.elsevier.com/locate/tecto Deep structure of the northeastern Japan arc and its implications for crustal deformation and shallow seismic activity Akira Hasegawa*, Junichi Nakajima, Norihito Umino, Satoshi Miura Research Center for Prediction of Earthquakes and Volcanic Eruptions, Graduate School of Science, Tohoku University, Sendai 980-8578, Japan Received 16 September 2004; received in revised form 18 March 2005; accepted 29 March 2005 Available online 10 May 2005 Abstract Seismic tomography studies in the northeastern Japan arc have revealed the existence of an inclined sheet-like seismic low- velocity and high-attenuation zone in the mantle wedge at depths shallower than about 150 km. This sheet-like low-velocity, high-attenuation zone is oriented sub-parallel to the subducted slab, and is considered to correspond to the upwelling flow portion of the subduction-induced convection. The low-velocity, high-attenuation zone reaches the Moho immediately beneath suggests that the volcanic front is formed by this hot upwelling flow. Aqueous fluids supplied by the subducted slab are probably transported upward through this upwelling flow to reach shallow levels beneath the Backbone Range where they are s a s n an s s crustal rocks, resulting in local contractive deformation and uplift along the Backbone Range under the compressional stress field of the volcanic arc. A strain-rate distribution map generated from GPS data reveals a notable concentration of east-west contraction along the Backbone Range, consistent with this interpretation. Shallow inland earthquakes are also concentrated in the upper crust of this locally large contraction deformation zone. Based on these observations, a simple model is proposed to explain the deformation pattern of the crust and the characteristic shallow seismic activity beneath the northeastern Japan arc. ① 2005 Elsevier B.V. All rights reserved. Keywords: Arc magmatism; Aqueous fluids; Crustal deformation; Shallow seismicity; Subduction zone; Northeastern Japan arc 1. Introduction mantle at a convergence rate of 8-9 cm/year and at an angle of about 30°. Many shallow earthquakes occur Northeastern Japan is located at a subduction zone, beneath the Pacific Ocean mainly along the upper where the Pacific plate subducts downward into the boundary of the Pacific plate associated with its subduction. Beneath the land area, shallow earth- quakes also occur in the upper crust; many of them are * Corresponding author. Tel.: +81 22 225 1950; fax: +81 22 264 concentrated in a long, narrow zone extending along 3292. the volcanic front or the central mountainous range E-mail address: hasegawa@aob.geophys.tohoku.ac.jp (A. Hasegawa). (Ou Backbone Range) which runs through the 0040-1951/S - see front matter @ 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2005.03.018 60 A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 middle of the land area nearly parallel to the trench earthquakes (Nagai et al., 2001; Yamanaka and axis (Fig. 1). Kikuchi, 2004; Matsuzawa et al., 2002; Okada et Great progress has been made in the last few al., 2003, Hasegawa et al., in press). years in understanding the stress concentration Understanding the mechanism of stress concen- mechanism causing interplate earthquakes beneath tration that leads to shallow inland earthquakes the Pacific Ocean off the northeastern Japan arc. (intraplate earthquakes) in the arc crust, on the other Asperities are distributed in patches surrounded by had, has advanced more slowly. Why, of the many stable sliding areas on the plate boundary. Aseismic active faults, does stress concentrate along just one of slip in the surrounding stable sliding areas results in them, leading to slip and an earthquake? It is to be the accumulation of stress at the asperities, and expected that once slip occurs on an active fault, earthquakes occur when the strength limit of an producing an earthquake, stress would become con- asperity is reached leading to sudden slip. It has centrated inregions adjacent to extensions of the fault, gradually become clear that this kind of asperity but in general, inland earthquakes occur in isolation, model (Lay and Kanamori, 1981） represents an and related earthquakes in adjacent regions are rarely accurate description of the mechanism of such if ever observed. Why is this so? Our current level of 138° 140° VF 142° 144° 42° 40° lwate Pref ref 38° 36° -10000-8000-6000-4000-2000 。 (m) Fig. 1. Map showing the northeastern Japan arc and its surroundings. Red triangles and thick gray line denote active volcanoes and the volcanic front, respectively. White arrow indicates the direction of the relative plate motion (Demets et al., 1994). The bathymetry is taken from the Japan Coast Guard. 1. Iwate volcano, 2. Naruko volcano. A. Hasegawa et al. / Tectonophysics 403 (2005) 59-75 61 understanding is not sufficient to explain these facts. It 2. Mantle wedge structure of the northeastern is clear that this scenario cannot be explained by a Japan arc simple model in which an elastic upper crust supports stress caused by relative plate motion, with slip (and Nakajima et al. (2001a,b), using data from the hence earthquakes) occurring when the stress exceeds seismic observation network, the density of which has the strength of the fault surface as a plane of weakness recently been increased, calculated the three-dimen- within the crust (Iio, 1996, 1998). sional seismic wave velocity structure for the north- Recent seismic tomography studies in the north- eastern Japan arc, updating the results of Zhao et al. eastern Japan arc have provided new information that (1992). Figs. 2 and 3 show the P-wave velocity (Vp) shows that water supplied by dehydration of the and S-wave velocity (Vs) on cross-sections perpen- subducting slab reaches the upper crust via the mantle dicular to the island arc. In any of the vertical cross- wedge, entrained in an upwelling flow in the mantle sections (a) to (f), the Pacific Plate subducting beneath that travels nearly parallel to the slab as a seismic low- the arc is imaged as a strong high-Vp and high-Vs velocity, high-attenuation zone in the mantle wedge. region. Within the mantle wedge immediately above The sheet-like upwelling flow aligned nearly parallel the Pacific Plate, low-Vp, low-Vs regions inclined to the slab reaches the Moho near the Backbone nearly parallel to the slab and extending from depths Range (or the volcanic front). Consequently, partial of about 100 to 150 km to the Moho appear clearly. melting is widely distributed along the volcanic front These regions of low seismic wave speed appear immediately below the Moho. When the molten clearly not only in cross-sections (a), (b), (d) and (f), material in such a melting zone approaches the which pass through active volcanoes, but also in surface, it cools and partially solidifies, expelling cross-sections (c) and (e), which do not include any water contained in the molten material. It is expected volcanoes. This illustrates the existence of a single that this water migrates to even shallower levels. sheet-like low-velocity zone inclined nearly parallel to Seismic tomography provides images of the upwelling the slab within the mantle wedge. This low-velocity paths of water in the upper crust as the low-velocity zone has high Vp/Vs values. Fig. 4 shows distribu- zones. The result is the continuous supply of water tion of Vp/Vs ratio at a depth of 40 km. We can see expelled from the subducting slab into a region below that a high Vp/Vs (and low Vp, Low Vs) zone is the Backbone Range. distributed along the volcanic front immediately Research on surface deformation based on GPS below the Moho. Similar low-velocity zones inclined data has revealed a zone of strain concentration that nearly parallel to slabs have also been observed in extends north-south along the Backbone Range, mantle wedges in other subduction zones (Abers, representing the local predominance of contractive 1994; Zhao et al., 1995, 1997; Gorbatov et al., 1999), deformation in the direction of relative plate motion although none are as clear as those in northeastern along the Backbone Range. This zone of strain Japan (Figs. 2 and 3). concentration is located above where the upwelling Seismic attenuation structure provides additional flow in the mantle wedge reaches the Moho. The information on the physical states of the earth's concentrated supply of water originating from the interior. Three-dimensional P-wave attenuation struc- slab must weaken the crustal material, causing ture beneath NE Japan was estimated by a joint contractive deformation to occur locally, that is, inversion for source parameters, site response and Qp anelastic deformation occurs locally even within the values (Tsumura et al., 2000). Fig. 5 shows across-arc upper crust. It is inferred that since this anelastic vertical cross-sections of Qp values along three lines deformation is non-uniform in space, shallow inland in the inserted map. Low Qp (high attenuation) zones earthquakes serve as a mechanism for making the inclined nearly parallel to the slab are clearly seen for overall deformation more uniform. Based on the all the cross-sections, although the extent of drop in present data, we propose this model of stress Qp-value is not large for cross-sections A and B. The concentration mechanism as a model for the occur- low-Qp zones are consistent with the inclined low-V rence of shallow inland earthquakes in the north- zone in Figs. 2 and 3. Thus there exists an inclined eastern Japan arc. sheet-like low Vp, low Vs, high Vp/Vs and low Qp 62 A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 Japan NEJapan Japan NEJapan Sea PacificOcear Sea PacificOcean 50 ept 100 150 a (d) 200 100 1 150 (b) (e) 200 60 100 150 200 dyp 6 3 Velocity perturbation (%) Fig. 2. Across-arc vertical cross-sections of P-wave velocity perturbations along lines in the inserted map of NE Japan (Nakajima et al., 2001a). The solid line and red triangles at the top represent land area and active volcanoes, respectively. Open and red circles denote earthquakes and deep, low-frequency microearthquakes, respectively. zone in the mantle wedge beneath the northeastern 1998), which is then dragged downward to a depth of Japan arc. 150-200 km where dehydration decomposition occurs (Iwamori, 1998; Schmidt and Poli, 1998). Slightly low velocity areas are imaged immediately above the 3. Upwelling flow within the mantle wedge subducted slab in some of vertical cross-sections of Figs. 2 and 3 (e.g., Fig. 3(b), (d), which might We infer that the inclined sheet-like low-V and correspond to this temporary layer of serpentine and low-Q zone described above corresponds to the chlorite, although more studies with much higher upwelling flow in the secondary convection (McKen- resolutions are required to confirm it. The water zie, 1969) accompanying slab subduction. Since released by this dehydration at depth is then trans- temperature increases with depth, the interior of this ported upward, encountering the upwelling flow at upwelling flow is at a higher temperature than the depths of 100-150 km. The supply of water to the surrounding region, and as such should have lower upwelling flow has the effect of lowering the solidus viscosity. In an old plate subduction zone such as temperature. From a comparison of the seismic wave northeastern Japan, water supplied from dehydration attenuation structure described in the previous section of the subducted slab may form a temporary layer of (Tsumura et al., 2000) with laboratory experiment serpentine and chlorite in the mantle wedge immedi- data, the temperature within the low-V, low-Q zone is ately above (Davies and Stevenson, 1992; Iwamori, estimated to be higher than that of the peridotite wet A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 63 Japan NEJapan Japan NEJapan Sea PacificOcear Sea PacificOcean 50 100 150 a (d) 200 100 50 (b) (e) 200 C 200 dys Velocityperturbation(%) Fig. 3. Across-arc vertical cross-sections of S-wave velocity perturbations along lines in the inserted map (Nakajima et al., 2001a). Other symbols are the same as in Fig. 2. solidus (Nakajima and Hasegawa,2003a). Further, the slab, was separated from the upper surface of the Nakajima et al., in press inferred from the ratio of fall- slab by about 50 km, and extended to depths of no off rates of P-wave and S-wave velocities that melt more than 125 km, accurately reproducing the low- inclusions are included in the low-V, low-Q zone, velocity zone observed in northeastern Japan (Figs. 2 having aspect ratios of 0.01-0.1 and volume fractions and 3). of 0.1 to several percent. The inferred water transport paths in the northeast- The existence of such a low-velocity zone inclined ern Japan subduction zone are shown schematically in nearly parallel to the slab at depths of less than 150 Fig. 6(a). The upwelling of hot mantle material from km, as detected by seismic tomography, has also been depth and the addition of water may cause partial confirmed by numerical simulation of the secondary melting with a volume fraction on the order of 0.1 to convection that accompanies plate subduction. Eberle several percent. Melt is formed both by decompres- et al. (2002) performed a numerical simulation of the sion melting and melting due to water addition. From corner flow that accompanies plate subduction using a the fact that the inclined low-velocity zone is only temperature-dependent viscosity coefficient, and clearly observed at depths shallower than about 150 found that a low-velocity zone with velocities several km (Zhao and Hasegawa, 1993), it is inferred that percent slower than in the surrounding region was melting by the addition of water plays an important generated, which would correspond to the present role in melt formation. Thus, water that originated upwelling region. The low-velocity zone determined from the slab is eventually incorporated into the melt. by Eberle et al. (2002) was aligned nearly parallel to The upwelling flow including this melt eventually 64 A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 139Y 140Y 141Y 142Y fast directions in the back-arc region are nearly parallel to the direction of relative plate motion. Most Depth = 40 km of stations with such trench-perpendicular directions 41Y are located above the inclined low-velocity zone (i.e. upwelling flow) in the mantle wedge. The observed trench-perpendicular fast directions would be explained by lattice preferred orientation of minerals caused by flow-induced strain in the mantle wedge (Ribe, 1992; Tommasi, 1998; Zhang and Karato, 40Y- tionmethod such as travel-time inversion because of are seen in the fore-arc region. Perhaps another 机 mechanism is working to cause these directions in timeshave a very complicated behaviour including reflected waves. Uncertainties of model parameters 39Y- important information on the variation of magma Japan suggesting the wide occurrence of velocity reversals undulations and significant jumps. This situation PacificOce 0 38Y- Depth(km) 50 100 150 200 37Y B 0 the computational instability arising from the high Vp/Vs ray-tracing technique (Iwasaki, 1988), in which the 50 1.651.701.751.80 ing head waves and diffracted waves from the edges 100 Fig. 4. Vp/Vs ratio at a depth of 40 km (Nakajima et al., 2001a). analysis, we did forward modelling based on a 2D 150- reaches the Moho immediately below the volcanic 200 front, resulting in the accumulation of large amounts C 0 of melt immediately below the Moho along the volcanic front. Seismic tomography clearly reveals 2002a,b). As shown in Fig. 3, the recorded travel- 50 this continuous distribution of partially molten mate- 100- rial along the volcanic front and immediately below the Moho as a region of low Vp, low Vs, high Vp/Vs 150 and low Qp (Figs. 2 through 5). From this point of view, the volcanic front can be regarded to form 200 where a sheet-like upwelling flow in the mantle 3 5 6 wedge reaches the Moho. (1/Qp)*1000 of layer boundaries as well as conventional diving and reflections from steeply dipping interfaces, whose ray paths are shown in (b). (b) Ray diagrams for shot L-6. (c) Travel-time plots of shot M-5. Fig. 5. Across-arc vertical cross-sections of P-wave attenuation structure along lines in the inserted map (Tsumura et al., 2000). Red seems to support the existence of this upwelling flow and blue colors represent high and low attenuations, respectively, in the mantle wedge. Fig. 7 clearly shows a systematic according to the scale at the bottom. Other symbols are the same as spatial variation in directions of fast shear-waves. The in Fig. 2. A. Hasegawa et al. / Tectonophysics 403 (2005) 59-75 65 Volcanoes JapanSea NEJapan PacificOcean (a) 0 令a.f.ormelt Moho low-Fevents diapir 1 en) (pa V (km) 100 Depth d ite V high Plate cific Pacit 200 0 100 200 300 AcrossArcDistance(km) (b) Quaternaryvolcanoes CE crust PacificPlate oceaniccrust low-V Zone (partially molten) Fig. 6. (a) Schematic diagram of vertical cross-section of the crust and upper mantle of NE Japan, showing the inferred transportation paths of aqueous fluids. (b) Schematic 3D structure of the crust and upper mantle of NE Japan showing the upwelling flow with varying thickness in the mantle wedge. formation along the island arc. Recently, Tamura et al. These cross-arc bands in which Quaternary volcanoes (2002） investigated the distribution of Quaternary are concentrated coincidewith regionsof elevated volcanoes in northeastern Japan, and found that the topography and low Bouguer gravity anomaly. volcanoes are distributed in long and narrow bands Tamura et al. (2002) concluded that volcanoes form perpendicular to the island arc, forming 10 clusters of where inclined hot fingers (upwelling regions) dis- volcanoes occupying an average width of 50 km. tributed across a width of 50 km in the mantle wedge 66 A. Hasegawa et al. / Tectonophysics 403 (2005) 59-75 140Y 140.5Y 141Y 141.5Y 142Y 40Y 39.5Y- 39Y 35 38.5Y- 0.21 0.14 0.07 -6 -3 0 3 6 Delay time (sec) Velocity perturbation (%) Fig. 7. Direction of fast shear-wave and delay time plotted at each station superposed on shear-wave velocity perturbations in the mantle wedge (Nakajima and Hasegawa, 2004). Black lines denote the direction of fast shear-wave and length is proportional to the average time delay between the leading and following shear-waves. Velocity image is the shear-wave velocity perturbations along the inclined low-velocity zone in the mantle wedge as in Fig. 8(a). Red triangles show active volcanoes. White arrow indicates the direction of the relative plate motion (Demets et al., 1994). at a depth of 150 km reach the surface. The repeated S-wave velocity perturbations taken along the inclined supply of magma from hot fingers in the mantle low-velocity zone. The value is that along the surface wedge to the crust immediately above causes the ofminimumS-wavevelocitywithinthemantle bedrock to be uplifted and Quaternary volcanoes to wedge, and thus the figure shows the distribution of form. They further concluded that the magma that is S-wave velocity perturbations along the curved sur- supplied accumulates beneath the Moho, producing face joining the core of the low-velocity zone. As the local low Bouguer gravity anomalies. Tamura et al. (2002) predicted, the extent of velocity To confirm the model of Tamura et al. (2002), we drop within the low-velocity zone varies clearly along attempt to image the low-velocity zone in the mantle the strike of the island arc. wedge with a higher spatial resolution (Hasegawa and Comparing these results with the topographic map Nakajima, 2004). In this study, the velocity structure (Fig. 8(b)), we can see that there is good agreement outside of the mantle wedge was fixed to that obtained between the regions where the velocity drop is locally earlier by Nakajima et al. (2001a), and the velocity particularly strong in the low-velocity zone distributed distribution within the mantle wedge was estimated from 30 to 150 km depth in the mantle wedge and the using the same data set. The spatial resolution was 10 regions where elevations in the topography are high km or finer in both the horizontal and depth from the Backbone Range to the back-arc region. directions. The distribution of S-wave velocity Quaternary volcanoes (red circles)are distributed in obtained is shown in Fig. 8(a). The figure shows the those regions. In addition, low-frequency microearth- A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 67 139° 140° 141° 142 139° 140° 141° 142° (a) (b) 41° dVs Topography 40° 39° 38° 37° 。low-F microearthquakes 6 3 6 ·Quaternaryvolcanoes 0 400800120016002000 Velocityperturbation(%) -activefaults Altitude (m) Fig. 8. (a) S-wave velocity perturbations along the inclined low-velocity zone in the mantle wedge of NE Japan. (b) Topography map of NE Japan. Deep low-frequency microearthquakes were located by the Japan Meteorological Agency and Okada and Hasegawa (2000). Thick lines denote active faults (Active Fault Research Group, 1991). quakes (white circles) produced at depths of 25-40 these observational facts,are shown schematically in km, believed to be caused by sudden movements of Fig. 6(b), which shows a three-dimensional expan- fluids in the crust (Hasegawa et al., 1991; Hasegawa sion of the two-dimensional cross-section in Fig. and Yamamoto, 1994), are seen to occur immediately above zones of particularly large velocity drop in the sheet-like, with a thickness that varies locally from mantle wedge. place to place, rather than occurring in fingers as Spatial correlations between the following features suggested by Tamura et al. (2002). The volcanic front can be clearly seen in Fig. 8: 1) Regional variation of is formed where this upwelling flow finally contacts low-velocity zone distributed from 30 to 150 km the Moho. As the flow approaches the Moho, it slows depth in the mantle wedge, 2) The distribution of down. The melt contained in the flow accumulates low-frequency microearthquakes occurring at 25-40 over a wide area along the volcanic front, immedi- km depth, 3) The distribution of Quaternary volca- ately below the Moho, resulting in the low-velocity, noes at the surface, 4) The distribution of topo- high-attenuation zone that is seen to extend over a graphical elevations extending from the Backbone wide areas along the volcanic front. Seismic tomog- Range toward the back-arc region. The structure of raphy has revealed that in the volcanic zones, the crust and upper mantle in northeastern Japan, and differentiation occurs and magma rises to the middle the upwelling flow in the mantle, as inferred based on crust (see Figs. 12 and 13). 68 A.Hasegawa et al./ Tectonophysics 403 (2005) 59-75 package includes one or two velocity reversals (Fig. analysis software. Using this feature, the coordinates the back-arc side where the sheet is locally thick and of an isolated observation point can be estimated from there is a large amount of melt, part of the melt the observation data for that point alone, without sometimes separates from the upwelling flow before it having to form a baseline. For this estimation, reaches the Moho along the volcanic front. According parameters estimated in advance to high precision to the estimate by Nakajima et al., in press obtained by JPL including orbital histories of GPS satellites, using the rates of decrease of Vp and Vs, the volume clock errors and the Earth's rotation are used. Data fractions of melt within these regions of the back-arc were analyzed using this Precise Point Positioning side in the upwelling flow are on the order of 0.1% to km/s at depths of 5-15 km. The depth of layer several percent. The separated melt rises straight East-west components of horizontal strain rates upward in the plumes, and accumulates beneath the estimated from observational data from January 1997 Moho. Part of it continues to rise and penetrates into to December 2001 are shown in Fig. 9. Constraints interpretation for our model, where geological infor- have been applied so as to ensure that the strain rate is rock. We infer that this is how the concentrations of continuous in space (Miura et al., 2002; Sato et al., Quaternary volcanoes and elevations of topography 2002). The east-west components are shown as the extending from the volcanic front toward the back-arc direction in which the deformation that accompanies region formed, and probably this formation process the plate convergence predominates; north-south continues today. The alternation of the regions where Quaternary volcanoes are concentrated and regions without volcanoes in the direction along the island arc is presumed to be due to the variation of partial 41Y melting in the upwelling flow in the mantle wedge at depths from 30 to 150 km along the island arc. 0 10-7/yr. 4. Zones of concentrated deformation along the velocity layer cannot be uniquely determined from 40Y Observational data on the surface deformation field obtained from the nationwide GPS continuous obser- vation network (GEONET) of the Geographical Survey Institute of Japan have provided a great deal 39Y of information that was previously impossible to obtain, such as that related to the temporal and spatial is characterized by a thick (5-8 km) sedimentary Suwa et al.(2003) and Sato et al.(2002) have analyzed data from the GEONET and the observa- tional network of Tohoku University from 1997 to 38r 2001 seeking to clarify surface deformation in the Tohoku region. GIPSY-OASIS II (GPS Inferred Positioning System-Orbit Analysis and Simulation Software II), developed by the Jet Propulsion Labo- ratory (JPL) of the American National Atmospheric 37Y and Space Administration (NASA) was used for GPS 139Y 140Y 141Y 142Y data analysis. This analysis software estimates param- Fig. 9. Distribution of horizontal east-west strain rate estimated from GPS observations for the period from 1997 to 2001 (Sato et probability variables without the need to take double al., 2003). Contour interval is 100 ppb/year. Red triangles denote phase differences. This is a big advantage over other (10-25 km). The estimation error for the reflector A. Hasegawa et al. / Tectonophysics 403 (2005) 59-75 On their westward extension, a rather high-velocity components are much smaller than east-west compo- nents. It can be seen from Fig. 9 that there is a belt- like zone in which contractive deformation is con- horizontal extent is limited (<10-20 km), probably centrated along the Backbone Range (or the volcanic front). The concentrated zone in which contractive deformation predominates in the direction of relative 10-7/yr. plate motion is therefore distributed in a long and narrow band that runs throughout Tohoku along the Backbone Range. 5. Deformation of the arc crust and shallow inland earthquakes-their relationship with fluids Fig. 8. Record section and ray diagrams for shot L-2. (a) Record section with calculated travel-times (solid lines). A reduction velocity of 6.0 As shown in Section 3,the inclined sheet-like upwelling flow in the mantle wedge reaches the Moho along the volcanic front, that is, the Backbone Range. The distribution of the Vp/Vs ratio immediately below the Moho is shown in Fig. 4. The upwelling flow, imaged as low-Vp, low-Vs, high-Vp/Vs and low-Qp regions, is distributed nearly continuously along the volcanic front, immediately below the Moho. The melt incorporated into the upwelling flow either butts up against the bottom of the crust or penetrates into the crust. When it cools in the crust km/s is employed. Phase identifications for R7 to R12 are the same as given in Fig. 7(b). (b) Ray diagrams showing interpreted reflection phases 139Y 140Y 141Y and partially solidifies, water is expelled from it and 142Y moves upward. Thus, water of slab origin is supplied Fig. 10. Distribution of horizontal east-west strain rate for the continuously to the shallow part of the crust along the period from 1997 to 2001 (Sato et al., 2003), and shallow Backbone Range. The presence of water is consistent earthquakes located by the seismic network of Tohoku University for the same period. with the concentration of low-frequency microearth- quakes (Hasegawa et al., 1991; Hasegawa and the lowermost crust (Hasegawa et al., 1991; Hase- Yamamoto, 1994) at depths near the Moho, and with gawa and Yamamoto, 1994). The water forms a sill at S-wave reflectors at intermediate crustal depths (Hori intermediate crustal depths and accumulates, perhaps et al., 2004) along the Backbone Range. The presence corresponding to the bright S-wave reflectors that of water can be expected to weaken the crustal have been detected across a wide area along the material and to produce local contractive deformation Backbone Range (Matsumoto and Hasegawa, 1996; under a compressive stress field. We infer that this Hori et al. 2004). In the Backbone Range, the happens in the concentrated deformation zone along temperature is locally increased by the infiltration of the Backbone Range as shown in Fig. 9. This high-temperature material from the upper mantle, and concentrated deformation zone is also the location of the bottom of the seismogenic layer (the boundary considerable present microearthquake activity, as between brittle and ductile layers) is locally elevated shown in Fig. 10. (Hasegawa and Yamamoto, 1994; Hasegawa et al., The deformation pattern of the arc crust in north- 2000). Corresponding to this, the observed crustal eastern Japan inferred from these observed facts is heat flow has locally high values in the Backbone schematically shown in Fig. 11(a). As melt cools and Range (Furukawa, 1993; Tanaka and Ishikawa, 2002). solidifies, water that have separated from the melt The water continues to rise and reaches the upper sometimes moves suddenly in the lower crust, and is crust, causing plastic deformation in some part of the observed as deep low-frequency microearthquakes in brittle upper crust. 70 A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 Backbone Range (a) contraction&uplift (partly anelasticdeformation) WEST small earthquakes EAST large earthquake seismogenic zone brittle to ductiletransition low-V S-wave reflectors lower crust Moho low-V low-F events low-Q upper mantle (b) Backbone Range partly anelastic elasticdeformation deformation elasticdeformation large contraction reverse reverse small contraction fault fault large contraction Fig. 11. (a) Schematic illustration of across-arc vertical cross-section of the crust and uppermost mantle, showing the deformation pattern of the crust and the characteristic shallow seismic activity beneath NE Japan. (b) Map view schematically showing the deformation pattern of the upper crust. In the Backbone Range, where the seismogenic formed in this way, although some extension is layer is locally thin and melt and water are observed locally around Iwate volcano, which is distributed in the lower crust, the entire crust will probably related to the volcanic activity of Mt. be locally weak in comparison with the surrounding Iwate, which started in 1998. Numerical simulation region. For this reason, the arc crust, which is being studies are essential to obtain a quantitative model compressed in the direction of relative plate motion, having spatial perturbationsof elastic and viscous deforms elastically outside of the Backbone Range, rheological constants which can explain the observed but anelastically in part within the upper crust along amount of the deformation, however, they are left for the Backbone Range, which can be expected to future studies. cause local contraction and uplift. We infer that the Research on surface deformation based on analysis concentrated deformation region shown in Fig. 9 was of GPS data (Sato et al., 2003) is steadily revealing A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 71 evidence of uplifting zones along the Backbone Prefecture and southern Iwate Prefecture, also has a Range as predicted by the present model. Local concentration of shallow microearthquakes (Fig. 10). contractive deformation along the Backbone Range This region includes the hypocenters of the 1900 is perhaps caused by asseismic slip on the deep Northern Miyagi earthquake (M7.0) and the 1962 extension of faults and/or by plastic volume deforma- Northern Miyagi earthquake (M6.5). Vp, Vs and Vp/ tion in the lower crust, leading to stress concentration Vs ratios in east-west vertical cross-section along a in the upper crust immediately above. Anelastic line across this region are shown in Fig. 12. In this deformation may also occur partially, in the upper region, a large amount of data is available from crust. This eventually leads to the rupture of the whole densely spaced temporary observation stations, mak- upper crust, producing a shallow inland earthquake ing imaging possible at higher spatial resolution that makes the deformation uniform in space (lio et (Nakajima and Hasegawa, 2003b). The cause of the al.，2000,2002).Anelastic contractive deformation concentrated deformation zone on the fore-arc side, along the Backbone Range including the upper crust which could not be understood from the image causes numerous shallow microearthquakes as it immediately below the Moho (Fig. 4), can perhaps advances, as seen in Fig. 10. be understood from Fig. 12. In addition to the low- Fig. 9 shows that, there is one more long, narrow velocity zone extending from below the Moho region on the fore-arc side where contractive defor- beneath the Backbone Range to immediately below mation predominates,in addition to the Backbone the Naruko volcano, there is another low-velocity Range where the upwelling flow in the mantle wedge zone that branches off and extends to the eastern side reaches the Moho. This region, in northern Miyagi (the fore-arc side). Naruko volcano Naruko volcano 140.25 E 141.25 E dv(%) 10 10 (km) 5 20 properties. According to the seismic reflection data, 30 Dep 40 (a) dvp b dvs -10 50 (Adachi, 2002; Iwasaki et al., 2003a,b). Such reflec- (km) 000 (km) 20 h10 00 Dept pth 30 Dep 40 tion patterns imply that the crust of the KA is (C) Vp/Vs our velocity structure, we interpret the low velocity 4.06 and 3.43 km/s, forming velocity reversals at the provided from vertical seismic profiling (VSP) data 30 Distance (km） 0○C 50 Vp/Vs bright S-wave reflectors Electrical Resistivity low-F microearthquakes 1.61 11.681.75 1.82 1.89 Fig. 12. EW vertical cross-sections of (a) P-wave and (b) S-wave velocity perturbations, (c) Vp/Vs (Nakajima and Hasegawa, 2003b), and (d) electrical resistivity (Mitsuhata et al., 2002) along line a in Fig. 10. Rectangles in (a), (b) and (c) show the range of cross-section in (d) in both horizontal and vertical directions. Red circles and dots denote low-frequency microearthquakes and shallow earthquakes, respectively. Red lines shallow earthquakes. 72 A. Hasegawa et al. / Tectonophysics 403 (2005) 59-75 Nakajima and Hasegawa (2003b) inferred from the Fig. 13 shows the Vp/Vs ratio on a vertical cross- rates of decrease of Vp and Vs that about 1% melting section along the Backbone Range. Regions of high occurs in the upper mantle, while several percent Vp/Vs ratio, believed to be regions of partial melting, melting occurs in the lower crust, with about 0.3-5% are distributed immediately beneath two volcanic water in the upper crust in this low-velocity zone. An areas, in the north and south, reaching intermediate MT survey conducted in the hypocentral region of the crustal depths. In these two areas, the amount of melt 1962 Northern Miyagi earthquake detected a clear supplied from the mantle wedge, and consequently the low-resistivity zone in almost exactly the same amount of water, must be greater than the area location as the low seismic velocity zone as shown between the two areas. Accordingly, within these in Fig. 12(d) (Mitsuhata et al., 2002). Immediately two areas, it can be expected that weakening of the above this zone there is a sheet-like distribution of crust will be considerable and that local contractive microearthquakes, dipping to the west, representing deformation will proceed rapidly. If this is the case, aftershocks of the 1962 Northern Miyagi earthquake then stress will be concentrated in the area between (Kono et al., 1993). these two areas, perhaps causing a reverse fault From these observations, we infer that water of earthquake to occur at the edge of the Backbone slab origin is supplied not only to the Backbone Range (or inside it), as shown schematically in Fig. Range but also to the hypocentral region of the 1962 11(b). In fact, the fault plane of some large earth- Northern Miyagi earthquake on the fore-arc side. It is quakes such as the 1896 Rikuu earthquake (M7.2) conceivable that this causes local contractive defor- was not within these two volcanic areas, but at the mation in this region as well as in the Backbone western and eastern edges of the area (or inside it) zone, Hokkaido (Japan), revealed by vibroseis seismic reflection between them (Active Fault Research Group, 1991). In the location where contractive deformation Even within the volcanic area, it appears that the occurs locally, not only are microearthquakes con- same kind of phenomenon is taking place, although centrated (Fig. 10), but large earthquakes that rupture on a smaller scale. Fig. 14 shows the S-wave velocity the entire seismogenic layer also occur. We infer that distribution at a depth of 4.5 km in the Onikobe area contractive deformation occurs principally as anelastic of northern Miyagi Prefecture (Onodera et al., 1998), deformation where the entire crust including the upper which is part of the above-mentioned volcanic area. In crust has been locally weakened. The observations this area, the lower boundary of the seismogenic layer given below suggest that since this kind of anelastic (the brittle to ductile transition zone) is relatively deformation does not proceed uniformly in space, shallow, on the order of 7 km (Hasegawa et al., 2000). large earthquakes that cause the overall contractive The estimated velocity distribution shows that the deformation tobecome uniform occur at locations of velocity within the caldera is low while that outside smaller contractive deformation. the caldera is high. It is expected that more water is Mt. Naruko Mt.Kurikoma 0 supporting plate tectonics in Hokkaido. Pap. Meteorol. Geo- foreland fold-and-thrust belt, Hokkaido, Japan. J. Japanese 10 20 1998. Crustal structure and tectonics of the Hidaka collision 30 The western part of the profile across the NJA, 50 50 100 0 150 Distance (km) Vp/Vs bright S-wave reflectors low-Fmicroearthquakes 1.611.681.751.82 (Ed.), Handbook of Physical Properties of Rocks, vol. 2. CRC Fig. 13. NS vertical cross-section of Vp/Vs structure in NE Japan along the line in the inserted map (Nakajima et al., 2001b). Other symbols are the same as in Fig. 12. A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 73 dVs 1996/8/11 M5.9 140.5°E 140.6°E 140.7°E 140.8°E Sanzugawa caldera Mt.Kurikoma 1996/8/11 M5.7 1985/3/28 M5.3 1996/8/13 M4.9 Onikobe caldera 38.8N 140° 142° Mukaimachi caldera Mt.Naruko 40° 38.7°N 1976/7/5 38 -10 0 10 M4.9 厂 5km Velocityperturbation(%) Fig. 14. S-wave velocity perturbations at 4.5 km depth (Onodera et al., 1998) and fault planes of earthquakes (Umino et al., 1998) in the Onikobe area shown in the inserted map. Fault planes of earthquakes with magnitudes greater than ~5 are indicated by rectangles. Arrows in each fault plane show slip vectors. Small circles denote aftershocks of the M5.9 Onikobe earthquake sequence in 1996. Caldera rims are indicated by bold lines (Yoshida, 2001), and red triangles denote active volcanoes. supplied within the caldera than outside the caldera, phenomenon is similar to that shown schematically in and consequently there will be considerable anelastic Fig. 11(b), but on a smaller scale. contractive deformation within the caldera. In 1996 there was considerable seismic activity in this region, with the largest earthquake a M5.9 event. In this 6. Concluding remarks region, where the seismogenic layer is locally thin, and at depths of around 7 km, the M5.9 earthquake In the northeastern Japan arc, shallow earthquakes was sufficient to rupture the entire seismogenic layer are concentrated in a region of large contractive (Umino and Hasegawa, 2002). From Fig. 14, rela- deformation in the direction of relative plate motion. tively large earthquakes for this region (M5 class) Research based on a comparison of crustal horizontal occur not inside the calderas, but around them. In deformation rates over the last 100 years has particular, the M5.9 earthquake occurred between the previously confirmed that such a region also corre- Sanzugawa caldera and the Onikobe caldera. Thus, sponds to a region of low seismic velocity (Hasegawa the M5.9 earthquake occurred in the region between et al., 2000). Based on these observations, Hasegawa the calderas to compensate for the delay in the et al. (2000) inferred the upwelling of water from progress of anelastic contractive deformation. This depth to weaken the crust and increase local crustal 74 A. Hasegawa et al. / Tectonophysics 403 (2005) 59-75 contraction rates, resulting in shallow crustal earth- active volcanoes in northeastern Japan. Tectonophysics 233, quakes in such areas. 233-252. Hasegawa, A., Nakajima, J., 2004. Geophysical constraints on slab In the present paper, this concept was extended, subduction and arc magmatism, AGU Geophys. Monograph, and a simple model was proposed based on the high- 150, IUGG volume 19, 81-94,2005. 209-229 (in Japanese with English abstract). Hasegawa, A., Zhao, D., Hori, S., Yamamoto, A., Horiuchi, S., mined by seismic tomography and detailed crustal 1991. Deep structure of the northeastern Japan arc and Hokkaido Univ., Ser. 7, Geophysics 10, 31-52. its relationship to seismic and volcanic activity. Nature 352, 683-689. Monogr. Assoc. Geol. Collab. Jpn. 31, 173-187. Hasegawa, A., Yamamoto, A., Umino, N., Miura, S., Horiuchi, S., tion in the island arc and the occurrence of shallow Zhao, D., Sato, H., 2000. Seismic activity and deformation inland earthquakes in northeastern Japan. Although process of the crust within the overriding plate in the north- the validity of this model must await future verifica- eastern Japan subduction zone. Tectonophysics 319, 225-239. Hasegawa, A., Uchida, N., Igarashi, T., Matsuzawa, T., Okada, T., 31, 189-203 (in Japanese with English abstract). Miura, S., Suwa, Y., in press. Asperities and quasi-static slips on deformation and in the occurrence of shallow inland the subducting plate boundary east off Tohoku, NE Japan, earthquakes. SEIZE volume, Columbia University Press. Hori, S., Umino, N., Kono, T., Hasegawa, A., 2004. Distinct S-wave reflectors (bright spots) extensively distributed in the crust and Acknowledgments upper mantle beneath the northeastern Japan arc. J. Seismol. Soc. Jpn., Ser. 2 (56), 435-446 (in Japanese). lio, Y., 1996. A possible generating process of the 1995 southern We wish to express our thanks for comments by H. Hyogo prefecture earthquake—stick of fault and slip on Iwamori, Y. Iio, T. Matsuzawa, and two anonymous detachment. J. Seismol. Soc. Jpn., Ser. 2 (49), 103-112 (in reviewers which have contributed importantly to Japanese). improve this paper. This work was partially supported lio, Y., 1998. ohmin—a role on the generation of earthquake. J. by a grant from the Ministry of Education, Culture, Seismol. Soc. Jpn., Ser. 2 (50), 273-281 (in Japanese). lio, Y., Kobayashi, Y., Sagiya, T., Shiozaki, I., 2000. Water in the Sports, Science and Technology of Japan. lower crust controls the deformation of the arc crust. Earth Mon. 247, 37-44 (in Japanese). lio, Y, Sagiya, T., Kobayashi, Y, Shiozaki, I., 2002. Water- References weakened lower crust and its role in the concentrated deformation in the Japanese Islands. Earth Planet. Sci. Lett. 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Hasegawa (2005) deep structure of northeastern japan.txt
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RESEARCH ARTICLE
Origin and evolution of the Paleo-Kuril arc inferred from
detrital zircon U –Pb chronology in eastern Hokkaido, NE Asia
Futoshi Nanayama1,2| Toru Yamasaki1| Toshiya Kanamatsu3|
Hideki Iwano4,5| Tohru Danhara4| Takafumi Hirata5
1Geological Survey of Japan, AIST, Tsukuba,
Japan
2Museum of Natural and Environmental
History, Shizuoka, Japan
3R&D Center for Earthquake and Tsunami,
JAMSTEC, Yokisuka, Japan
4Kyoto Fission-Track Co., Ltd., Kyoto, Japan
5Geochemical Research Center, Graduate
School of Science, The University of Tokyo,
Tokyo, Japan
Correspondence
Futoshi Nanayama, Geological Survey of
Japan, AIST, 1-1-1Higashi, Tsukuba, Ibaraki
305-8567, Japan.
Email: nanayama-f@aist.go.jp
Funding informationJSPS KAKENHI, Grant/Award Numbers:
JP19K04025, JP21H04523Abstract
The Nemuro and Saroma Groups and Yusenkyo Formation occur in eastern Hokkaido
and are considered to be forearc or intra-arc basin sediments of the Paleo-Kuril arc
(PKA) deposited during the Late Cretaceous to middle Eocene. To further clarify the
origin of the PKA, we examined the U –Pb ages of detrital zircons within these sand-
stones and acidic tuff beds; based on our results, we drew the following conclusions.(1) The PKA originated from an oceanic island arc on the oceanic Izanagi Plate around
85 Ma, to which the Nikoro Group greenstone complex was accreted between
81–80 Ma; the Lowest Unit of the Saroma Group covered the surface of the Nikoro
accretional greenstone complex. (2) The PKA then first collided with NE Asia around
the beginning of the deposition of the Hamanaka Formation (~70 Ma) and transi-tioned to a continental arc adjacent to NE Asia. This collision generated giant slump
beds during deposition of the Akkeshi Formation. During deposition of the Kiritappu
Formation, the entire Nemuro Peninsula was uplifted, supplying large volumes of
clastic sediments. (3) Although we do not have data directly bearing on why the
North American Plate was established in the edge of NE Asia, we speculate that it
separated from the Eurasian continent around ~70 Ma when NE Asia first collided
with the PKA. Subsequently, the PKA has behaved as part of the North AmericanPlate. The Paleo-Japan arc (or East Sikhote –Alin arc) and the PKA became joined via
the Hidaka Belt. Around 40 Ma, during the deposition of coarse conglomerate beds
of the Urahoro Group, the PKA was uplifted and bent clockwise due to a second col-
lision with NE Asia. (4) The modern Kuril arc was established around 36 Ma (late
Eocene –early Oligocene), and the Kuril backarc basin opened into the present tec-
tonic setting in the late Oligocene –early Miocene.
KEYWORDS
collision, LA-ICP-MS, late cretaceous, NE Asia, origin, Paleo-Kuril arc, zircon U –Pb age
1|INTRODUCTION
The Kuril –Kamchatka arc (hereafter, simply “the Kuril arc ”) extends
along the southern coast of the Kuril Islands from the southern part ofthe Kamchatka Peninsula in the northwestern Pacific Ocean to the
southeastern part of Hokkaido, Japan (Figure
1). The Kuril arc –trenchsystem extends further south to the Japan Trench and further north
to the southwestern coast of Bering Island, where it orthogonally
intersects the Aleutian Trench. The Kuril arc results from the subduc-
tion of the Pacific Plate beneath the Okhotsk microplate as part of theNorth American Plate (Seno et al.,
1996 ) in the Kuril subduction zone
(Kuril Trench). In our opinion below, we do not use the Okhotsk PlateReceived: 10 February 2022 Revised: 30 July 2022 Accepted: 9 August 2022
DOI: 10.1111/iar.12458
Island Arc. 2022;31:e12458. wileyonlinelibrary.com/journal/iar © 2022 John Wiley & Sons Australia, Ltd. 1o f3 2
https://doi.org/10.1111/iar.12458
OH
GYEurasia Plate
Japan TrenchPacific PlateKuril Arc
(Kuril-Kamchatka Arc)Aleutian Trench
01 2 k mAmur microplateNOR
Kuril Basin
Japan
BasinNorth Asian
Craton
Sino-Korean
CratonKuril Trench (Kuril-Kamchatka Trench) East Sikhote-Alin Volcanic BeltOkhotsk-Chukotka
Volcanic Belt
PhilippineSea PlateOkhotsk microplateBering SeaN72°
E168°N64°
N48°
N40°
E144°
E132°E120° E108°North American
Plate
Back arc basin
Strongly thinning
continental crust
Weakly thinning
continental crustNTOAT
TAT
SYHKOM
N56°
Cratons (Archean and Proterozoic: 4-0.5 Ga)
Cratonal margins (Riphean, Cambrian,
Devonian through Jurassic: 1.6 GB-145 Ma)
Cratonal terranes and Superterranes (Archean through Jurassic: 4 Ga-145 Ma)
East Sikhote-Alin and Okhotsk-Chukotka volcanic belts(Early Cretaceous through Paleogene: 145-50 Ma)
Accretional complex and forearc basin sediments around
Sakhalin-Hokkaido and western Kamchatka
(Late Jurassic-Paleocene)
Accretied palaeo-arcs aroud Sea of Okahotsk
(Late Cretaceous-Paleocene: 85-41.2 Ma) WK
Sredinny MassifKvakhona Arc TerraneWK
Koriak
FIGURE 1 Legend on next page.2o f3 2 NANAYAMA ET AL.
14401738, 2022, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/iar.12458 by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
as the Okhotsk microcontinent, but use the Okhotsk microplate as
part of the North American Plate. However, if the Okhotsk Plate isused in the original paper, both terms will be listed together.
Granodiorites of the Prasolov magmatic complexes are exposed
around Tyatya Volcano in the northeastern part of Kunashir Island.Although these rocks were considered to result from Mesozoic mag-
matism (Nemoto et al.,
1940 ) along the Paleo-Kuril arc (PKA). De
Grave et al. ( 2016 ) reported a zircon U –Pb age of 31.3 ± 1.1 Ma, dat-
ing the origin of the modern Kuril arc to as late as the early Oligocene
(De Grave et al., 2016 ; Vaes et al., 2019 ; Figure 2). K–Ar dating has
revealed contemporaneous igneous activity in eastern Hokkaido,including the Futamata andesite lava (31.4 ± 0.7 Ma; augite –hyper-
sthene –hornblende andesite) and the Nuibetsu Formation (32.6
± 1.7 Ma; augite andesite agglomerate) (Shibata & Tanai,
1982 ).
Recently, Yabe et al. ( 2021 ) reported a new K –Ar age of 36.6
± 2.0 Ma for the Futamata andesite in the Tsubetsu area. These
results constrain early PKA magmatism to the late Eocene to early Oli-
gocene (Figures 2–4).
Since the 1970s, several researchers have proposed that the geo-
logical bodies partially exposed in eastern Hokkaido were derived
directly from the PKA and the related Paleo-Kuril arc –trench system
(Kimura & Tamaki, 1985 ; Nanayama et al., 1993 ; Okada, 1982 ,1983 ;
Sakakibara et al., 1986 ), or from an undefined arc –trench system pre-
sumed to extend north –south along the western edge of the Okhotsk
Plate (Okhotsk microcontinent; Kiminami, 1986 ; Kiminami &
Kontani, 1983a ). Other researchers have a contrasting view of the
geological history around the Sea of Okhotsk. For example, ( i) Zharov
(2005 ) described the above-mentioned paleoarc, the Tokoro arc in
eastern Hokkaido, and the Patience arc (Terpeniya island arc) of the
Terpeniya peninsula on the eastern margin of Sakhalin Island, Russian
Far East, and ( ii) Hourigan et al. ( 2009 ) described the Olyutorsky arc
terrane of central Kamchatka. These two paleoarcs formed in the Late
Cretaceous and were thrust upon NE Asia in the Eocene (Vaes
et al., 2019 ; Zharov, 2005 ; Figure 1). Accordingly, and based on seis-
mic tomography, paleomagnetic data, and geological reviews, Domeieret al. (
2017 ) and Vaes et al. ( 2019 ) interpreted that paleo-island arcs
comprising the Olyutorsky paleoarc, the East Sakhalin –Nemuro
paleoarc (equivalent to the PKA), and the Kronotsky paleoarc collidedwith the Eurasian continent during the Eocene.
Considering the transitional tectonic setting of the modern and
Paleo-Kuril arc, it is difficult to conclude whether the PKA generatedthe geological members of NE Asia. Furthermore, since the 1980s, no
geological information has offered an explanation of how and when
the geological members of eastern Hokkaido were incorporated intothe North American Plate.In this study, we used laser ablation inductively coupled plasma
mass spectrometry (LA-ICP-MS) to obtain the U –Pb ages of detrital
zircon grains extracted from sandstones and acidic tuff beds of the
Nemuro Group (Nokkamappu, Monshizu, Hamanaka, Akkeshi, and
Kawaruppu Formations) and the Saroma Group (the Lowest Unit andYusenkyo Formation), which were part of the Paleo-Kuril arc –trench
system. Our results provide important new constraints on the origin
of the Paleo-Kuril arc –trench system.
2|OVERVIEW OF THE PALEO-KURIL
ARC –TRENCH AND ITS TECTONIC SETTING
2.1 |Origin of the Okhotsk paleoland and the
Paleo-Kuril arc –trench system
Based on geophysical and structural data showing that most of the
Sea of Okhotsk area consists of continental crust comprising a micro-continent (Figure
1), Den and Hotta ( 1973 ) recognized the area
including the Sea of Okhotsk, the central and eastern parts of
Hokkaido, Sakhalin Island, Kamchatka Peninsula, and the Koriak areaas the Okhotsk Plate (Okhotsk microcontinent). Based on the move-
ment of this area, they regarded the plate as an independent minor
plate. Dickinson (
1978 ) and Parfenov et al. ( 1981 ) considered that the
collision between the northern edge of the Okhotsk Plate (Okhotskmicrocontinent) and NE Asia caused the new arc –trench system to
jump to the southern edge of the plate. This hypothesis centers on
the existence of “an extended granitic horizon ”under the Sea of
Okhotsk in geophysical data (Maruyama & Seno,
1986 ; Moskvina
et al., 1983 ; Tuyezov, 1971 ).
Kiminami et al. ( 1978 ), Kiminami and Kontani ( 1979 ), and Kontani
and Kiminami ( 1980 ) studied paleocurrents recorded in turbidite sand-
stones in the Nemuro, Saroma, and Yubetsu Groups along the island
arc in the eastern Sea of Okhotsk, and suggested that an “Okhotsk
paleoland ”once existed. Furthermore, Kontani and Kiminami ( 1980 ),
Kiminami ( 1983 ), and Kiminami and Kontani ( 1983b ) suggested that
the Okhotsk paleoland was a magmatic arc based on the sandstone
compositions there. However, they did not consider that this paleoarcwas distinct from the PKA, and was instead another independent arc
along the western edge of the Okhotsk Plate (Okhotsk microconti-
nent). In addition, the hypothesized Okhotsk paleoland was inferredfrom strata distributed in eastern Hokkaido; the above studies did not
report the assumed area, volume, or age of the paleoland, and these
geological data have yet to be estimated. Importantly, the hypothesisof the Okhotsk paleoland was based mainly on the idea that the
FIGURE 1 Geodynamical summary of NE Asia. The map is derived from a generalized Northeast Asia geodynamics map at 1:10 000 000
scale (Parfenov et al., 2011 ). The basement of the Sea of Okhotsk was interpreted from seismic data by Xu et al. ( 2016 ). NOR, north Okhotsk rise;
GY, Gyeonggi –Yeongnam Cratonal terranes (Archean –Proterozoic, 4 –0.5 Ga); KOM, Kolyma –Omolon Jurassic Superterrane (Archean –Jurassic,
4 Ga to 145.0 Ma); OH, Okhotsk Cratonal terrane; SYH, Sorachi –Yezo and Hidaka belts (late Jurassic –early Eocene, 152.1 –50 Ma); WK, West
Kamchatka terrane (middle cretaceous –paleocene, 125.0 –56.0 Ma); TAT, Terpeniya arc terrane (late cretaceous –paleocene, 85 –56.0 Ma); OAT,
Olyutorsky arc terrane (late cretaceous –paleocene, 85 –56.0 Ma); NT, Nemuro and Tokoro belts (late cretaceous –middle Eocene, 85 –41 Ma)NANAYAMA ET AL. 3o f3 2
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Pacific OceanNemuro
Belt
Kuril TrenchVityaz Ridge
Shikotan Is.
140°E 145°E50°N
45°N
150°E 135°EEast Sikhote-Alin Arc Terrane (Early Cretaceouce -Paleocene)
400 kmJapan Basin
(Late Oligocene ~)Sea of Okhotsk
East Sakhaline TerraneSchmidt ophiolite complex
Academy of Sciences
Plateau
Kuril Basin
(Late Oligocene ~)
Kuril Arc (Late Eocene - Early Oligocene ~)Tonin-Aniva
Complex
Sorachi
-YezoBeltWest
Sakhaline Terrane
Rebun
-Kabato
BeltTerpenia Arc
TerraneEurasia
HidakaBelt
Tokoro
BeltKm IdSamarka Terrane
(Jurrasic - Early Cretaceous accretionalcomplex)N
Zhuravlevka Terrane (Early Cretaceouce )
OshimaOshima
BeltBeltOshimaBelt
Sea of Japan
FIGURE 2 Geological map of the Hokkaido –Sakhalin Islands and the eastern Eurasian continent (Sikhote –Alin) since the early
cretaceous. See Zharov ( 2005 ), Ueda ( 2016 ), Vaes et al. ( 2019 ), and Nanayama et al. ( 2021 ) for geological informat i o no nS a k h a l i nI s l a n d
and the Sikhote –Alin area in NE Asia. Km, Kamuikotan metamorphic zone; ID, Idonnappu zone. The red triangular mark indicates the
locality where De Grave et al. ( 2016 ) reported the zircon U –Pb age of 30.9 ± 0.8 Ma in the Prasolov magmatic complexes on Kunashir
Island. The orange triangular mark indicates the locality where Yabe et al. ( 2021 )r e p o r t e dt h eK –Ar age of 36.6 ± 2.0 Ma in the Futamata
andesite. The yellow triangular mark indicates the locality where Shibata and Tanai ( 1982 )r e p o r t e dt h eK –A ra g eo f3 2 . 6±1 . 7M ai nt h e
Nuibetsu formation4o f3 2 NANAYAMA ET AL.
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Okhotsk Plate originated from a microcontinent (Maruyama &
Seno, 1986 ).
In contrast, Kiminami ( 1986 ) attributed the origin of the Paleo-
Kuril arc –trench system to the collision of the Okhotsk Plate (Okhotsk
microcontinent) and Eurasian Plates along the northern margin of the
Sea of Okhotsk in the Koriak area (NE Asia) during the Late Creta-ceous (ca. 90 –80 Ma). The new arc –trench system then shifted to the
southern edge of the Okhotsk Plate (Okhotsk microcontinent), to
which Kimura (
1994 ) and Kimura ( 1996 ) attributed the supply of large
volumes of clastic sediments to the seafloor between NE Asia and the
Okhotsk microcontinent and areas further south, and the later forma-
tion of the Hidaka Belt collisional complex. However, that hypothesisno longer holds because the depositional age of the Hidaka Super-
group has been revised to Late Cretaceous –early Eocene (Nanayama
et al.,
2018 ,2019 ,2021 ).
Geological and geophysics investigations proposed geodynamic
models for the Sea of Okhotsk area and the Kuril Arc in the late 20 cen-
tury (e.g., Kiminami, 1986 ; Kimura, 1994 ,1996 ; Maruyama et al., 1997 )
and in the 21 centuries (Domeier et al., 2017 ; Kimura et al., 2019 ;S e t o n
et al., 2015 ; Vaes et al., 2019 ; Zharov, 2005 ;Z h a oe ta l . , 2018 ), but they
have not reached to consensus yet. Particularly, the eastern edge of NEAsia, including the Okhotsk microplate, is generally considered to belong
to the North American Plate (e.g., Domeier et al., 2017 ;S e n o
et al., 1996 ; Vaes et al., 2019 ) because the modern Kuril arc is behaving
as part of the North American Plate (e.g., Kimura & Tamaki, 1985 ;V a e s
et al., 2019 )( F i g u r e 1). On the other hand, the Okhotsk microplate or
Okhotsk Plate established due to a clockwise rotation around the Sea ofOkhotsk region at 30 Ma related to the collision of the Indian subconti-
nent at 40 Ma (e.g., Kimura & Tamaki,
1985 ,1986 ). In addition, there is
no evidence that the Okhotsk microcontinent (Okhotsk Plate) hasexisted before 30 Ma as bellow.
2.2 |Geological structure of the three zones of
Hokkaido associated with the PKA
According to Ueda ( 2016 ) and Nanayama et al. ( 1993 ,2021 ), the
basement in the Paleo-Kuril arc –trench system and eastern Hokkaido
can be divided into three zones from west to east: the Hidaka Belt,
the Tokoro Belt, and the Nemuro Belt (Figures 2and3).
The Nemuro Belt is exposed in two areas, the Kushiro –Nemuro
(Konsen) coast and the Shiranuka hills (Figure 3), and comprises the
Doto magnetic
anomaly belt?Numuro BeltTokoro
Belt
Hidaka
Belt
Nakanogawa
GroupYubetsu
Group
Yusenkyo
Fm.
Konsen coast S
Shiranuka
hills
Toyokoro
hillsKitami
hills
Nemuro
HamanakaKushiro
AkkeshiAbashiriHama
saroma
Kitami
Obihiro
Daimaruyama and
Tachiiwa greenstonesNorthern
Unit
Sourthern
UnitSYusenkyo Fm.:
140.0 ± 8.7 °
Nakanogawa Gr.
(Southern Unit):
174.0 ± 25.2 °Tokoro Belt
Hidaka Belt
NNNemuro Gr.:
0.1 ± 23.3 °
Nemuro Gr.:
-20.8 ± 14.5 °Nemuro Belt
(Konsen coastalarea)
Nemuro Belt (Shiranuka hills area)
N NNemuro Group: 73.2 ± 7.2 °
Late Cretaceous - Middle EoceneUrahoro Group: 38.1 ± 22.7 °
Middle EoceneHidaka Supergroup
Nemuro Group
Saroma GroupUrahoro Group, Rikubetsu and
Wakamatsuzawa formations
Nikoro GroupNN
50 km
Tokoro metamorphic rocks43°N144°N 143°N 144°N
44°N
FIGURE 3 Distribution of terranes belonging to the Paleo-Kuril arc (modified from Katagiri et al., 2019 ) and paleomagnetic declinations of the
terranes with 95% confidence limits (Fujiwara et al., 1995 ; Fujiwara & Kanamatsu, 1990 ; Hamano et al., 1986 ; Nifuku et al., 2009 ). The Yusenkyo
formation was deposited during the late Campanian –early Maastrichtian (Kanamatsu et al., 1992 ) and the Nakanogawa group during the
Paleocene –early Eocene (Nanayama et al., 2019 ). Gr, group, Fm, formationNANAYAMA ET AL. 5o f3 2
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Black mudstone
and sandstonePelagic acidic tuff and mudstone
Limestone Chert and
pelagic mudstone
Conglomerate
40
50
60
70
80
90
100
110
120
130
140
150
160
170
180
190
201
252
299PALEOGENE CRETACEOUS JURASSIC TRI.
EARLY MIDDLE LATE ELAP YLRAE LATE EOCENE3020OLIGNEO
MIOCEPER.13
Volcanic rocks56–54
58–57Southern Hidaka belt
Northern
UnitSouthern
Unit
65–6358–575349–48 52–49 K –Ar
?Tokoro belt
Nikoro &
Saroma Grs.Yubetsu Group
77 6749–48
Exotic blocksNakanogawa Group
60–57 K–Ar
Saroma FaultNorthern Hidaka belt
ca. 90 K ArErimo Fm.Spreading stage of the Kuril basinSpreading stage of the Kuril basin
Early volcanic stage Early volcanic stage
of the Kuril arcof the Kuril arc
Bend and uplift stageBend and uplift stage
?Nemuro belt
Nemuro Gr.
(Shiranuka hills)
Urahoro Gr.
40–39Abashiri Tectonic LineWakamatsuzawa
and Rikubetsu Fms.
Upper Unit
Middle Unit
Lower Unit Katsuhira
Fm.Kawaru
-ppu Fm.31–30
Lowest Unit
Yusenkyo Fm.?Nemuro belt
Nemuro Gr.
(Konsen coast)
Urahoro Gr.
40–3929–26 FTTsubatsu &
Kawakami Grs.
Kiritappu Fm.
Tokotan Fm.
Akkeshi Fm. (Upper,
Middle, Lower Mems.)
Hamanaka Fm.
Monshizu Fm.
Ohtamura Fm.Nokkamappu Fm.Furubanya Unit
Konbumori Fm.48 FT
76 K-Ar70 K-Ar66 K-Ar 67 K-Ar54 K-Ar39-34 K-Ar34-31 K-ArPrasolov magmatic
complexes
Futamata And.
Nuibetsu Fm.
Shiomi Fm.
FIGURE 4 Geochronological correlation chart comparing sedimentary rocks in the Tokoro and Nemuro belts and the sedimentary complexes
(Nakanogawa and Yubetsu groups) east of the Hidaka Belt in Central –Eastern Hokkaido. Modified after Ueda ( 2016 ) and Nanayama et al. ( 2018 ,
2019 ,2021 ); Nanayama, Kurita, et al. ( 2020 ); Nanayama, Watanabe, et al. ( 2020 ). The red number +K–Ar shown the reported K –Ar ages (Ma),
red number +FT shown the reported fission track ages (Ma) and the black number indicates the reported zircon U –Pb ages (Ma) of acidic tuff,
volcanic rocks and sedimentary rocks in the eastern Hokkaido. Gr(s), group(s), Fm(s), formation(s)6o f3 2 NANAYAMA ET AL.
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Nemuro Group, which consists of Upper Cretaceous to middle Eocene
strata and various igneous rocks (basalt to andesite and granitic rocks).Kimura and Tamaki (
1985 ) reconstructed the Nemuro Group to have
been situated farther north before the extension of the Kuril backarc
basin. Other igneous rocks make up a remnant arc extending from thesouthwestern tip of the Kamchatka Peninsula toward the Academy of
Science Plateau on the north side of the Kuril backarc basin. These
two groups of igneous rocks originally formed the crust of the PKAand its forearc basin (Kiminami,
1983 ; Kimura & Tamaki, 1985 ;
Okada, 1982 ,1983 ; Figure 2).
The eastern margin of the Nemuro Group along the Konsen coast
can be traced to the Habomai and Shikotan Islands (Lesser KurilIslands; Hirano et al.,
2012 ) and even further to the Vityaz Ridge
(Lelikov & Emelyanova, 2011 ; Figure 2). Dredge samples of the Vityaz
Ridge, as well as the on-land Nemuro Belt, indicate that the modernKuril arc formed upon the old igneous crust of the PKA (Lelikov &
Emelyanova,
2011 ).
In the Shiranuka hills, the Nemuro Group strikes NNE –SSW, and
Kiminami ( 1986 ) traced its northern extension to the Kotikovaya
Group (Tosy et al., 2005 ) on the Terpeniya Peninsula, eastern Sakhalin
Island (Figure 2). In particular, the Uchir Formation at the bottom of
the Kotikovaya Group consists of volcanic rocks, volcanic sandstone,
and mudstone, lithologically similar to the Nokkamappu Formation at
the bottom of the Nemuro Group. Sasa and Nishida ( 1937 ) reported
Inoceramus schmidti (middle Campanian) in the lower part of the Koti-
kovaya Group, and Tosy et al. ( 2005 ) reported its depositional age to
be Maastrichtian –Eocene based on fossil radiolarians, pollen, and
plants. The depositional age and lithology of the Kotikovaya Groupare thus very similar to those of the Nemuro Group. Based on paleo-
magnetic data, Nanayama et al. (
1993 ,2021 ) interpreted that the
Nemuro Group in the Shiranuka hills was bent clockwise by 90/C14or
more from an initial nearly east –west strike (Figure 3; see also figure
4 in Nanayama et al., 1993 ), consistent with the Kotikovaya Group
being the northern extension of the Nemuro Group. Furthermore, theTerpeniya Peninsula extends southeast to the bottom of the Sea ofOkhotsk and is cut by the Kuril backarc basin; closing the backarc
basin by section restoration shows that the Terpeniya Peninsula was
originally nearly continuous with the Nemuro Group in the Shiranukahills (Figure
2).
The Tokoro Belt consists mainly of the Nikoro Group, a Late Cre-
taceous accretionary greenstone complex (Sakai et al., 2019 ;
Sakakibara et al., 1986 ,1993 ; Yamasaki & Nanayama, 2017 ) associ-
ated with the low-temperature, high-pressure Tokoro metamorphic
rocks that occur along its eastern margin (Sakakibara, 1991 ; Figures 3,
4). This accretionary greenstone complex formed a structural high,
which the Saroma Group covered to form the forearc basin associated
with the Nemuro Group (Nanayama et al., 1993 ; Sakakibara
et al., 1986 ). Its southern extension, exposed through small fensters
to the west of the Shiranuka hills, is thought to continue to the Toyo-
koro hills (Figure 3). The extension of the boundary between the
Tokoro and Nemuro belts has been clearly traced to the Doto mag-netic anomaly in Eastern Hokkaido, which suggests the subsidence of
a large serpentinite body (Morijiri & Nakagawa,
2014 ; Figure 3).The Urahoro Group, consisting of alluvial-fan and fluvial-
conglomerate beds associated with shallow-marine sediments, alsooccurs in the Tokoro Belt. The dominant cobble to boulder clasts in
this group are red cherts and greenstones derived from the Nikoro
Group, and detrital chromian spinel occurs in the Urahoro Groupsandstone. It is thus highly probable that the Nikoro Group and ultra-
mafic rocks were exposed between off-Kushiro and Hamanaka
(Nanayama et al.,
1994 ; Figure 3).
The Hidaka Belt is an accretionary or collisional complex mainly
composed of large volumes of clastic rocks associated with in situ Late
Cretaceous –early Eocene greenstones (mid-ocean ridge basalts); these
clastic rocks were distributed beyond the trench area and consist ofdeep-sea turbidites (Nanayama et al.,
1993 ,2021 ). The Yubetsu and
Nakanogawa Groups, occurring on the eastern side of the Hidaka Belt
(Nanayama et al., 1993 ,2021 ; Sakakibara et al., 1986 ; Figure 3)w e r e
deposited in a large deep-sea fan complex between the PKA and the
Paleo-Japan or east Sikhote –Alin continental arc on the eastern margin
of NE Asia. Large volumes of clastic Hidaka Belt sediments derived fromboth the Paleo-Japan arc and the PKA mixed with the Yubetsu and
Nakagawa Group sediments (Nanayama et al.,
1993 ,2021 ).
The current strikes of the Tokoro Belt (Nikoro and Saroma
Groups) and the Yubetsu Group between the Shiranuka and Kitami
hills are NNE –SSW to NE –SW, as is the strike of the Nakanogawa
Group on the eastern side of the Hidaka Mountains. Paleomagneticdata obtained from the Nemuro Group in the Shiranuka hills and fromthe Yusenkyo Formation show that the Tokoro and Nemuro Belts
were bent clockwise by 90
/C14or more from an initial east –west strike in
the original stage of the Kuril arc –trench system (Hamano et al., 1986 ;
Kimura, 1990 ; Fujiwara & Kanamatsu, 1990 ; Kanamatsu et al., 1992 ;
Fujiwara et al., 1995 ; Katagiri et al., 2019 ; Figure 3). Correcting for
this bending requires that the PKA associated with the Olyutorsky arc(OAT in Figure
1) also originally had an east –west to ENE –WSW
strike (Hamano et al., 1986 ; Kimura, 1990 ; Kanamatsu et al., 1992 ;
Katagiri et al., 2019 ; Kimura & Tamaki, 1985 ; Nanayama et al., 1993 ,
2021 ; Vaes et al., 2019 ; Figure 3).
The Paleo-Kuril arc –trench system and PKA are thought to have
been active since the Late Cretaceous (Campanian, ca. 80 Ma). The
PKA may have originated as a continental arc along the southern edgeof the Okhotsk Plate (Okhotsk microcontinent) (Kiminami,
1986 ;
Kimura & Tamaki, 1985 ) or as an island arc on an oceanic plate, per-
haps the ancient Izanagi Plate (Kimura et al., 2019 ; Vaes et al., 2019 ).
However, to date, no compelling data have allowed the PKA's exact
origin to be distinguished.
2.3 |Terpeniya and Olyutorsky arc terranes in Far
East Russia
The geology of central Sakhalin Island correlates with that in central
Hokkaido and is roughly divided into three zones from west to east: the
West Sakhalin terrane, associated with the Kamuikotan –Susunai meta-
morphic zone (western accretional complex and forearc basin sediments
corresponding to the Sorachi –Yezo Belt); the East Sakhalin terraneNANAYAMA ET AL. 7o f3 2
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(corresponding to the Hidaka Belt); and the Terpeniya arc terrane (TAT;
eastern paleoarc and forearc basin sediments; Rozhdestvensky, 1986 ;Z h a o
et al., 2019 ;Z h a r o v , 2005 ;F i g u r e s 1and2).
The eastern forearc basin sediments comprise upper Campanian –
Eocene marine and fluvial sediments associated with volcanic rocks,called the “Terpeniya arc terrane (TAT) ”around the Terpeniya Penin-
sula and along the eastern margin of Sakhalin Island
(Melankholina,
1976 ; Sasa & Nishida, 1937 ; Figure 2) and the “Kotiko-
vaya Group ”by Tosy et al. ( 2005 ). These sediments are accompanied
by volcanic (basalt, andesite, and alkaline basaltic trachyandesite) and
ultramafic rocks (ophiolite). This ophiolite is continuous to the Schmidtophiolite complex at the northern edge of Sakhalin Island (Vaeset al.,
2019 ; Figure 2) because the pyroclastic sandstones of the
Schmidt ophiolite complex are accompanied by basaltic volcanic frag-
ments (Kameda et al., 2000 ). Rozhdestvensky ( 1986 ) estimated that
the Terpeniya arc formed as the different arc-trench system from the
East Sikhote –Alin arc during the Late Cretaceous (Figure 2).
Kiminami ( 1986 ) proposed that the TAT is the northern extension
of the Nemuro and Tokoro belts. Nanayama et al. ( 1993 ,2021 ) con-
sidered that the TAT belongs to the western margin of the PKA
because of the structural continuity prior to the opening of the Kurilbackarc basin (see Section
2.2). In contrast, other researchers have
considered the TAT to be a paleoarc that was thrust upon NE Asia
during the early Eocene (Zharov, 2005 ).
The Olyutorsky arc terrane (OAT) is a similar paleoarc fragment present
in central Kamchatka (Hourigan et al., 2009 ;K o n s t a n t i n o v s k a y a , 2001 ;
Figure 1). Vaes et al. ( 2019 ) argued that two distinct island arcs occurred in
the Late Cretaceous and were thrust upon NE Asia during the Eocene. TheWest Kamchatka terrane (WK; Figure
1) is a complex of various geological
bodies, the details of which are not yet clear. For example, the Sredinny
massif was thought to be the oldest ge ological body on Kamchatka Penin-
sula, comprising Devonian high-grade metamorphic rocks (Shul'diner
et al., 1979 ), but recently reported ages are limited to the Late Cretaceous
(Hourigan et al., 2009 ; Konstantinovskaya, 2001 ). In contrast, the Kvakhona
arc terrane on the west coast of Kamchatka Peninsula is of Middle Jurassicto Early Cretaceous age (Figure
1), and may, in fact, be the oldest unit there
(Konstantinovskaya, 2001 ). Importantly, the West Kamchatka terrane is
continuous with the Sea of Okhotsk seafloor to the west (Xu et al., 2016 ).
3|GEOLOGY AND DEPOSITIONAL AGES
OF THE NEMURO AND SAROMA GROUPS
Here, the strata, lithology, and depositional ages of each exposure are
reviewed.
3.1 |Nemuro Group in the Nemuro Belt
The Nemuro Group occurs in two areas, the Konsen coast and the
Shiranuka hills, and there are some lithological and structural differ-ences between the two areas (e.g., Kiminami,
1983 ; Figure 3).3.1.1 | Nemuro Group on the Konsen coast
On the Konsen coast, the Nemuro Group is at least 3000 m thick. It
generally strikes east –west to northeast –southwest and dips gently to
the south. Here, the strata and lithology of the Nemuro Group on theKonsen coast are summarized (Figure
4).
Nokkamappu Formation
The Nokkamappu Formation is typically distributed on the northern
coast of Nemuro Peninsula. This formation comprises tuff breccia, vol-
canic conglomerate, volcanic sandstone, tuffaceous siltstone, and pil-low lava. The igneous rocks associated with this formation arediverse, including not only alkaline basalts but also calc-alkaline andes-
ites to dacites (Ikeda & Goto,
2019 ). The presence of many large
bivalve fossils such as Inoceramus indicates that the depositional envi-
ronment was a shallow-marine area around active volcanoes
(Kiminami, 1983 ).
Two fossil ages have been reported for this formation: early –
middle Campanian to early Maastrichtian based on Inoceramus
schmidti (Matsumoto, 1970 ) and middle –late Campanian based on
nannofossils (Okada et al., 1987 ). Recently, Shigeta and Tsutsumi
(2018 ) reported the zircon U –Pb age of the 1.0-m-thick acidic tuff
interbedded in the Sphenoceramus (Inoceramus) schmidti Zone of the
Yezo Group to be 80.2 ± 0.8 Ma (middle Campanian).
The composition of this formation is typical volcanic sandstone,
containing prominent basaltic to andesitic rock fragments associated
with clinopyroxene and hornblende grains. The quartz content is
extremely low, and feldspar and rock fragments dominate; the rock isclassified as a feldspar to lithic wacke defined by Okada (
1971 ) on the
Q–F–R diagram (Kiminami, 1979 ; Figure 5).
Otamura Formation
The Otamura Formation is mainly composed of hemipelagic mud-
stone and alternating beds of hemipelagic mudstone and thinly bed-ded turbidite sandstone. This formation is intruded by a monzonite,which has been dated to 70.7 ± 2.2 Ma (biotite K –Ar age) to 69.5
± 1.4 Ma (biotite –alkali feldspar –whole-rock Rb –Sr isochron age)
(Shibata,
1986 ).
Monshizu Formation
The Monshizu Formation mainly comprises coarse-grained turbiditesandstones. These thickly stratified sandstones are accompanied by
sandstone-dominant alternating beds of turbidite sandstone and
hemipelagic mudstone. The Monshizu Formation is often accompa-nied by vitreous acidic tuff beds in certain horizons. Of these, the
3.0-m-thick Mon-AT acidic tuff in the uppermost horizon of this for-
mation is a useful marker bed.
Inoceramus and other fossils presumed to be Maastrichtian
(Matsumoto,
1970 ) or early Maastrichtian (Naruse et al., 2000 ).
The composition of the turbidite sandstone of the Monshizu
Formation is very similar to that of the Nokkamappu Formation(Kiminami,
1979 ;F i g u r e 5).8o f3 2 NANAYAMA ET AL.
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Oborogawa Formation
The Oborogawa Formation mainly comprises hemipelagic mudstone
associated with thinly bedded turbidite sandstone. Its depositional
age is estimated to be early Maastrichtian (Naruse et al., 2000 ). How-
ever, Shibata ( 1986 ) reported ages of 69.8 ± 1.7 Ma (biotite –whole-
rock Rb –Sr isochron age) for an intrusive monzonite, 69.0 ± 2.1 Ma
(whole-rock K –Ar age) for a dolerite sill, and 69.9 ± 2.2 Ma (biotite K –
Ar age) for another intrusive monzonite in the Oborogawa Formation.We consider the Rb –Sr age to be slightly younger than the fossil age
of Naruse et al. (
2000 ).
Hamanaka Formation
The Hamanaka Formation consists of sandstone-dominated alternat-
ing beds of turbidite sandstone and hemipelagic mudstone. Severaldolerite sheets (10 –30 m thick) are interbedded in this formation. The
6.3-m-thick Ham-AT acidic tuff bed in the lower horizon of this for-
mation is a useful marker bed. Naruse et al. (
2000 ) estimated the fossil
age of the Hamanaka Formation to be early Maastrichtian. Alkali feld-spar K –Ar ages of 65.5 ± 2.1 and 67.7 ± 2.1 Ma have been reported
for the intrusive dolerite sheets; these K –Ar ages are concordant but
slightly younger than the fossil age (Shibata,
1986 ).
According to Kiminami ( 1979 ), the Hamanaka Formation is volca-
nic sandstone mainly containing basalt to andesite fragments (about
90% of all rock fragments). However, the sandstone compositions ofthis formation are slightly more quartz-dominated than those of the
Nokkamappu to Monshizu Formations. This sandstone is classified as
a feldspar wacke (Figure
5). The heavy mineral assemblage consists of
clinopyroxene, hornblende, and small amounts of apatite, orthopyrox-ene, epidote, and zircon.
Akkeshi Formation
The Akkeshi Formation mainly comprises thick-bedded turbidite sand-
stone and alternating beds of turbidite sandstone and hemipelagicmudstone and is characterized by the occurrence of variously sized
slump beds in many horizons. Kiminami ( 1979 ) divided it into lower,
middle, and upper members based on differences of lithology and sed-
imentary character. The fossil age of Ammonite fossils in the lower
and middle member is estimated to be early Maastrichtian(Matsumoto & Yoshida,
1979 ; Naruse et al., 2000 ). On the other hand,
Okada et al. ( 1987 ) reported nannofossils of early paleocene (Danian)
age in the upper member. Therefore, the Cretaceous –Paleogene (K –
Pg) boundary (66 Ma) is presumed to be between the middle and
upper members of the Akkeshi Formation (Naruse et al., 2000 ).
According to Kiminami ( 1979 ), the sandstones of the lower and
upper members are rich in quartz (Q), with quartz/feldspar ratios
often exceeding 1.6; they are classified as lithic wacke or feldspar
wacke. Most grains of the upper member are rounded, and are thuspresumed to be more mature than those in the sandstone of the lowermember. Rock fragments in these members include chert, granite, and
schist, in addition to volcanic rocks (Figure
5). The main heavy min-
erals present are garnet, zircon, epidote, and apatite. In contrast, thesandstone of the middle member contains little quartz and is mainly
composed of feldspar (F) and rock fragments (R); it is classified as a
feldspar to lithic wacke. The heavy mineral assemblage and composi-tion is very similar to those of the sandstone in the lower member; it
contains abundant clinopyroxene, hornblende, and andesitic to dacitic
volcanic rock fragments (Kiminami,
1979 ).
Senposhi Formation
The Senposhi Formation is about 1300 m thick and is distributedalong the western coast of Akkeshi Bay; it is mainly composed ofhemipelagic mudstone associated with thinly bedded turbidite sand-
stones. Naruse et al. (
2000 ) reported an early Maastrichtian Ammonite
fossil and Okada et al. ( 1987 ) reported a latest Maastrichtian nanno-
fossil in this formation. However, Nifuku et al. ( 2009 ) correlated four
magnetic zones with polar ages C31r –C30n in the geomagneticPVAEMAEMADMADMA
RMAQ
F R50 % 50 %
50 %Saroma Group
Middle to Lower units
Lowset Uni t
Yusenkyo Fm.Nemuro Group ʢKonsen coast)
Tokotan Fm.
Akkeshi Fm. (Upper Member)Akkeshi Fm. (Middle Member)Akkeshi Fm. (Lower Member)Hamanaka Fm.
Monshizu Fm.
Ohtamura Fm.
Nokkamappu Fm.
Nemuro Group
ʢShiranuka hills)
Kawaruppu Fm.
Katsuhira Fm.FIGURE 5 Modal compositions of sandstones
in the Nemuro and Saroma groups and theirconstituent rock fragments. Sandstonecompositions are presented on the quartz(Q)–feldspar (F) –rock fragments (R) ternary
diagram and compared to fields discriminating arcprovenance (Kumon et al.,
2000 ): PVA, primitive
volcanic arc; EMA, evolved/mature magmatic arc;DMA, dissected magmatic arc; RMA, renewed
magmatic arc. Data for the Nemuro group are
from Kiminami (
1979 ) and Kontani et al. ( 1986 ),
and data for the Saroma group are from Kiminamiand Kontani (
1983b ). Fm, formationNANAYAMA ET AL. 9o f3 2
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record; thus, the depositional age most likely extends from middle to
late Maastrichtian (ca. 69 –67 Ma).
Tokotan Formation
The Tokotan Formation is mainly composed of hemipelagic mudstone
and associated conglomerate beds and thin-bedded turbidite sand-
stone, and is partly associated with the conglomerate bed of the
Kiritappu Formation (see below). Okada et al. ( 1987 ) reported early
Paleocene (Danian) nannofossils in this formation.
The sandstone of the Tokotan Formation contains little quartz
and is mainly composed of feldspar and rock fragments; it is classifiedas a feldspar to lithic wacke. Heavy minerals and rock fragments aregenerally clinopyroxene, hornblende, and andesitic to dacitic volcanic
rock fragments (Kiminami,
1979 ; Figure 5).
Kiritappu Formation
The Kiritappu Formation mainly consists of a coarse conglomerate
bed containing rounded to subrounded clasts of dolerite, basalt,andesite, granite, sandstone, and mudstone 3 –35 cm in diameter. The
total thickness of this formation is 300 m or more. Because this for-
mation shows the largest clast diameter of the Nemuro Group, it isinferred that the source area was uplifted and eroded during the
deposition of this formation (Kiminami,
1979 ). Yoshida ( 1967 )
reported planktonic foraminiferal fossils of middle to early Paleocene(Danian) age, and Okada et al. (
1987 ) reported contemporaneous
nannofossils.
Stratigraphy in the upper part of the Nemuro Group to the west of
Akkeshi Bay
In the western Konsen coastal area bounded by Akkeshi Bay, upper
Nemuro Group layers overlie the Kiritappu Formation, such as theShiomi and Konbumori Formations and the Furubanya unit which are
mainly composed of bioturbated silty sandstone associated with con-
glomerate and sandstone beds deposited on the continental shelf(Okada et al.,
1987 ) (Figure 4). They are presented here from oldest to
youngest. Late paleocene nannofossils have been reported in the
Shiomi Formation. Late paleocene (Selandian) nannofossils have been
reported in the Konbumori Formation. Early eocene (Ypresian) nanno-fossils have been reported in the Furubanya unit.
3.1.2 | Nemuro Group in the Shiranuka hills
The Nemuro Group is narrowly distributed in the Shiranuka hills. Thewestern edge of this group is cut by the ATL (Figure
3). The strati-
graphic thickness of the Nemuro Group in the Shiranuka hills is at
least 4000 m. The strata generally strike NNE –SSW and dip to the
east. The Nemuro Group in this area is divided into the Katsuhira andKawaruppu Formations (Kaiho,
1984 ).
Katsuhira Formation
The Katsuhira Formation mainly comprises hemipelagic mudstone
associated with thin-bedded turbidite sandstone beds deposited onthe continental shelf. The depositional age of this formation lies
between late Maastrichtian and the middle Paleocene (Selandian)(Kaiho,
1984 ).
According to Kontani et al. ( 1986 ), the composition of the sand-
stone is lithic wacke, containing ~20% quartz on the Q –F–R diagram
(Figure 5). Rock fragments mainly include sedimentary and volcanic
rocks with some minor granitic rocks. According to Nakazoe ( 1963 ),
the heavy mineral assemblage comprises garnet, epidote, and zirconwith minor clinopyroxene and hornblende. Most zircon grains are
columnar and euhedral; subrounded zircons are purple in color.
Kawaruppu Formation
The Kawaruppu Formation is mainly composed of continental-shelfsandstone and hemipelagic mudstone beds, and thickly bedded turbi-
dite sandstone or conglomerate beds are interbedded in four horizons(SC1–SC4), including the basal bed (SC1).
The depositional age of the Kawaruppu Formation is from the
middle Paleocene (early Selandian) to the early Eocene (early Lutetian)(Kaiho,
1984 ). A biotite K –Ar age of 54.0 ± 1.7 Ma (Shibata, 1984 )
and two fission-track ages of 52 and 47.7 Ma (Kimura & Tsuji, 1990 )
have been reported for acidic tuff layers directly above the SC3 hori-zon (Figure
4).
According to Kontani et al. ( 1986 ), the sandstone in this formation
contains little quartz (~10% on the Q –F–R diagram) and is a lithic
wacke (Figure 5). Rock fragments in this formation are dominated by
basaltic to andesitic volcanics. The heavy-mineral assemblage
reported by Nakazoe ( 1963 ) is similar to that of the Katsuhira Forma-
tion but in different proportions, consisting primarily of clinopyroxeneand hornblende with minor chromian spinel, garnet, zircon, and
epidote.
3.2 |Saroma Group in the Tokoro Belt
The Saroma Group is distributed in two narrow strips within the
Nikoro Group (Figure 3) and mainly comprises thick-bedded turbidite
sandstone, alternating beds of turbidite sandstone and hemipelagic
mudstone, and a thick hemipelagic mudstone associated with a thickbasal conglomerate bed. The stratigraphic thickness of this group is at
least 1300 m. From bottom to top, the Saroma Group is subdivided
into the lowest, lower, middle, and upper units (Research Group of theTokoro Belt,
1984 ).
The Lowest Unit is a thick conglomerate bed deposited by a sedi-
ment gravity flow. This unit has a maximum thickness of 100 m andunconformably covers the Nikoro Group. The clast assemblage is
characterized by subalkaline to alkaline volcanic (basalt, andesite, tra-
chyte, rhyolite) and plutonic rocks (syenite, monzonite) (ResearchGroup of the Tokoro Belt,
1984 ). The lower to upper units of the Sar-
oma Group are distributed around the Hamasaroma area (Figure 3).
These units consist of alternating beds of turbidite sandstone and
hemipelagic mudstone associated with large slump beds.
Sakakibara et al. ( 1986 ) and Nanayama et al. ( 1993 ) considered
the depositional environment of the Saroma Group to have been the10 of 32 NANAYAMA ET AL.
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forearc basin or slope basin overlying the accreted greenstone
complex of the Nikoro Group. The Research Group of the Tokoro Belt(
1984 ) clearly showed that the abundant andesite and alkaline plu-
tonic clasts of the Lowest Unit were derived from the PKA.
Information on the depositional ages of the Lowest Unit is lim-
ited. Sakakibara and Tanaka ( 1986 ) reported fragments of Inoceramus
sp. very similar to I. schmidti around Nikura River. U –Pb dating of
detrital zircons by Murakami et al. ( 2016 ) revealed that the youngest
cluster in the Middle and Upper Units is of Maastrichtian age
(72–66 Ma).
Sandstones in the Saroma Group are lithic wackes (partially are-
nites), poor in quartz and rich in volcanic rock fragments (basalt –
andesite) and clinopyroxene and hornblende grains.
3.3 |Yusenkyo Formation in the Tokoro Belt
The Yusenkyo Formation is distributed in a narrow fenster on the
western side of the Shiranuka hills and is considered to belong to the
Tokoro Belt (Kanamatsu et al., 1992 ; Figure 3). The Yusenkyo Forma-
tion is composed of volcanic sandstone, volcanic conglomerate, andtuffaceous mudstone. This ≥600-m-thick formation shows a mono-
cline structure. Late Campanian –early Maastrichtian radiolarian fossils
have been reported in the Yusenkyo Formation (Kanamatsuet al.,
1992 ). This formation also yields fragments of Inoceramus fossils
similar to those in the Lowest Unit of the Saroma Group (Kanamatsu
et al., 1992 ).
The sandstone of the Yusenkyo Formation is a quartz-poor lithic
wacke containing abundant greenish to reddish basalt and andesite
fragments (Figure 5). The heavy-mineral assemblage includes only
clinopyroxene and hornblende.
4|METHODS
We measured the zircon U –Pb ages of clastic rocks (e.g., turbidite
sandstones and acidic tuffs) surrounding the Nemuro and Saroma
Groups and the Yusenkyo Formation to re-examine their deposi-tional ages and better constrain their provenances. Detailed sam-
ple locations and descriptions are provided in Figures
6and 7.
Bulk samples weighed between 0.2 and 1.1 kg. Zircon grains wereseparated from sandstone and acidic tuff samples by grinding,
sieving, panning in water, magnetic separation, and heavy-liquid
separation. For zircon-bearing sa mples (seven sites, see Sections
5
and 6), batches of at most 100 zircon grains were embedded in
Perfluoroalkoxy (PFA) Teflon sheets and polished with diamond
paste. Cathodoluminescence images of the zircon grains in themount were acquired before U –Pb dating (Figure
8). Then,
from each site, we analyzed 25 highly euhedral and well-zoned
zircon grains clearly lacking any detrital core, as well as five
additional grains selected regardless of grain shape and structure.LA-ICP-MS analyses were performed on the core portions of the
selected grains.For most zircons (except those from sample ST-PS54-01, see next
paragraph),
238U–206Pb and235U–207Pb ages were determined from
Laser Ablation Inductively Coupled Plasma Mass Spectrometry
(LA-ICP-MS) analyses, performed using a CARBIDE femtosecond laser
ablation system (LIGHT CONVERSION, Vilnius, Lithuania, USA) and amagnetic sector multi-collector ICP-MS (Nu Plasma II; Nu Instruments,
Wrexham, UK). Helium was used as the carrier gas inside the ablation
cell, and it was mixed with argon before entering the ICP-MS. Single-spot laser ablation was performed with fluences of 3.7 –3.8 J cm
/C02.
The pit diameter was 10 μm. Three full-size electron multipliers were
used to measure202Hg,204(Hg+Pb), and208Pb simultaneously. The
206Pb,207Pb, and235Uo r238U signal intensities were monitored by
three Daly collectors, and232Th was monitored by a Faraday cup. The
collectors were described and characterized by Hattori et al. ( 2017 )
and Obayashi et al. ( 2017 ). The instrumentation and operating condi-
tions are summarized in Table S1. The mass bias on the207Pb/206Pb
and206Pb/238U ratios was externally corrected by replicate analyses
of the Nancy 91 500 (Wiedenbeck et al., 1995 ) or Ple ˇsovice (Sláma
et al., 2008 ) standards without correction for common Pb. The
207Pb/235U ratio was calculated based on the above two ratios and
assuming a constant238U/235U value of 137.88 (Jaffey et al., 1971 ).
Errors were calculated from the reproducibility of the primary zircon
standard and the counting statistics of the signal intensity of each iso-
tope. The quality of U –Pb analyses was confirmed by duplicate ana-
lyses of secondary standards GJ-1 (Jackson et al., 2004 ) and OD-3
(Iwano et al., 2013 ).
Sample ST-PS54-01 was analyzed with another LA-ICP-MS sys-
tem using a New Wave Research NWR-193 ArF laser ablation system(Fremont, California, USA; e.g., Sakata et al.,
2014 ) and a Thermo
Fisher Scientific iCAP-Qc quadrupole ICP-MS (Yokohama, Japan)
(Table S1). Helium was used as the carrier gas inside the ablation cell,and it was mixed with argon before entering the ICP-MS. The pit
diameter was 20 μm. Analytical conditions used for this sample are
reported in Table
1. Signal intensities for202Hg,204Pb (204Hg),206Pb,
207Pb,208Pb,232Th, and238U were obtained for 30 zircon crystals
from the sample. The Nancy 91500 and OD-3 zircon standards were
used as primary and secondary standards, respectively.
Based on the obtained ages, we selected the youngest grain pop-
ulation per sample as follows: (1) we selected grains with reliable and
concordant ages; (2) we calculated the weighted average of those
grain ages (Taylor, 1982 ); and (3) we selected the206Pb/238U age for
grains younger than 1 Ga and the207Pb/206Pb age for older grains.
The weighted average ages (YGC3 σ; young cluster 3sigma method
according to the definition of Coutts et al., 2019 ) were taken as the
formation age for each sample (2 σerrors).
5|SAMPLING POINTS FOR ZIRCON U –PB
DATING AND THEIR STRATIGRAPHIC
HORIZONS
Table 1, Figures 6, 7and S1 show the sampling points and horizons of
10 samples such as NM-NKK-01, Mon-AT-01, Ham-AT-01, KS-HMt-NANAYAMA ET AL. 11 of 32
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300 mNK-SRb-01
N44
0`25.03z E143
52`58.38z
(b) (a)
(d) (c)
(f) (e)
(g)HN-YS2-03
N43
10`26.29z
E143
42`58.46zKT-SRb-01
N43
54`10.43z
E143
48`522.90z
ST-PS54-01
N43
10`26.29z
E143
42`58.46z
ST-KWt-01
N43
10`26.29z
E143
42`58.46zAcidic tuff
54.0 ± 1.7 K-Ar
(Shibata, 1984)
Urahoro Gr.
Rushin
Fm. Yubetsu
Fm.
Mon-AT-01
N43
8`57.02z
E145
10`31.66zHam-AT-01
N43
9`10.28z
E145
14`19.42z
NM-NKK-01
N43
22`29.74z
E145
38`19.69zKS-HMt-01
N43
8`56.38z
E145
13`57.00zKS-AKb-01
N43
8`52.66z
E143
13`44.95z600 m
600 m
200 m400 m400 m400 mDiscovered locations
of Inoceramus sp.
(Sakakibara and Tanaka, 1986)
DoleriteOborogawa
Fm.Monshizu
Fm.Oborogawa
Fm.
Hamanaka
Fm.
Hamanaka
Fm.
Akkeshi
Fm.Hamanaka
Fm.Nikoro
Gr.Saroma
Gr.
Discovered locations of
radiolarian fossils indicated
late Campanian - early
Maastrichtian (Kanamatsu et al., 1992)
Discovered locations of
nannofossils indicated
middle - late Campanian
(CC19-CC22) (Okada et al., 1987)Nokkamappu
Fm.
Ohtamura
Fm.DoleriteDolerite
Nokkamappu
Fm.Nemuro Gr.
Kawaruppu Fm.Onbetsu Gr.F
Yusenkyo Fm.
Yusenkyo Fm.
Ponporoto coastMouraito
coast
Mouraito
PortHonbetsu
RiverKutonnikoro RiverNikura
River
Satonbetsu
River
FIGURE 6 Detailed sampling locations in the Nemuro and Saroma groups, eastern Hokkaido (see Figure S1 for broader geological context):
(a) NK-Srb-01; (b) KT-Srb-01; (c) HN-YS2-03 (d) ST-PS54 –01 and ST-KWt-01; (e) Mon-AT-01; (f) ham-AT-01, KS-HMt-01, and KS-AKb-01; and
(g) NM-NKK-01. Maps were adapted from a digital topographic map (scale 1:25 000) published by the geospatial information Authority of Japan12 of 32 NANAYAMA ET AL.
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01, KS-AKb-01, HN-YS2-03, KT-SRb-01, NK-SRb-01, ST-PS54-01,
and ST-KWt-01 used for zircon U –Pb dating in this study.
6|U–PB DATING RESULTS
The results of our analyses are reported in Tables S2 –S8 and plotted
in Figures 8–10.
6.1 |NM-NKK-01
We did not obtain any zircon grains from the 0.20 kg of sandstone
crushed.6.2 |Mon-AT-01
We obtained 250 zircon grains from 0.20 kg of vitreous tuff. This
sample yielded light brown, homogeneous, columnal, euhedral zircons
(Figure 8). All 30 analyzed grains were concordant (Figure 9) and seem
to form a single age cluster ranging from 72.60 ± 2.26 Ma to 67.81
± 3.17 Ma. We calculated the weighted average age (YGC3 σ)t ob e
69.59 ± 0.33 Ma (Figure 11).
6.3 |Ham-AT-01
We obtained 600 zircon grains from 0.20 kg of vitreous tuff. This
sample yielded light brown, homogeneous, columnal, euhedral zircons.
˒˒
˒
˒
NK-SRb-01Ham-ATMon-AT
HN-YS2-03
NM-NKK-01KS-HMt-01KS-AKb-01ST-PS54-01
ST-KWt-01Urahoro
Group(a) (b) (c)
(d) (e) (f)
(g) (h)(i)Dolerite
sheetSlump bed
˒
˒˒
˒˒˒
Ham-AT-01Mon-AT-01
FIGURE 7 Photographs of sampling sites in the Nemuro and Saroma groups, eastern Hokkaido (localities shown in Figures S1 and 6). (a) Sandy
mudstone sample ST-KWt-01 from the top of the Kawaruppu formation, covered by the Urahoro group along the Satonbetsu River, Shiranuka hills.(b) Sandstone sample ST-PS54-01 from the thick SC3 sandstone of the Kawaruppu formation along the Satonbetsu River, Shiranuka hills.(c) Sandstone sample KS-Akb-01 from a slump bed in the lower member of the Akkeshi formation along the Konsen coast. (d) Sandstone sample KS-HMt-01 from a thin turbidite bed in the upper horizon of the Hamanaka formation along the Konsen coast. (e) Vitric acidic tuff sample Ham-AT-01from the thick acidic tuff bed Ham-AT in the basal horizon of the Hamanaka formation along the Konsen coast. (f) Vitric acidic tuff sample Mon-AT-01 from the thick acidic tuff bed Mon-AT in the basal horizon of the Monshizu formation along the Konsen coast. (g) Sandstone sample NM-NKK-01from a thick shallow-marine sandstone bed in the Nokkamappu formation along the Nemuro coast. (h) Sandstone sample HN-YS2-03 from a thickgravelly sandstone bed in the Yusenkyo formation along the Konsen coast. (i) Sandstone sample NK-SRb-01 from a thick turbidite sandstone bed inthe basal formation of the Saroma Group in the Kitami hills. White arrows point stratigraphically upward for each bedNANAYAMA ET AL. 13 of 32
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Ham-AT-01 no. 2 ST-PS54-01 no. 3Mon-AT-01 no. 24Mon-AT-01 no. 1
NK-SRb-01 no. 14NK-SRb-01 no. 3
KS-HMt-01 no. 5 KS-AKb-01 no. 142.0 Ga1.8 Ga64.8 ʶ.B
68.9 ʶ.B82.6 ʶ.B 72.6 ± 2.3 Ma
69.5 ± 1.7 Ma
79.9 ʶ.B
FIGURE 8 Cathodoluminescence images of selected zircon grains from six sample SEM (scanning electron microscope) images of selected
zircon grains from two sample. U –Pb analytical spots (circles) and obtained ages are indicated (206Pb/238U ages for grains younger than 1 Ga and
207Pb/206Pb ages for older grains; see Section 4)14 of 32 NANAYAMA ET AL.
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TABLE 1 Data table of sample number, sampling horizon, location and field description and observation under the microscope
Sample number Sampling horizon Location Field description Observation under the microscope
NM-NKK-01 Nokkamappu
Formation, NemuroGroup on theKonsen coastThe western coast of
Cape Nokkamappu inMakinouchi, Nemuro(43
/C1422029.2100N,
145/C1438019.7100E;
Figure 6g)This sample was taken from outcrops of medium- to
coarse-grained sandstone beds 10 cm thick withinthe alternating sandstone and mudstone beds. Theseblack sandstones contain abundant andesite
fragments, and many bivalve fragments and outsized
volcanic clasts are scattered throughout. Parallellaminae and low-angle hummocky cross-stratification indicate that these sandstone bedswere deposited in a shallow-marine environment
(Figure
7g)This wacke sandstone contain many angular particles and have the
modal composition 0.0% quartz, 26.0% feldspar, 56.4% rockfragments, 2.1% heavy minerals, and 15.5% matrix. Feldspars aremostly anhedral to subhedral plagioclase, commonly showing albite
twins and sericitization. Rock fragments include basaltic volcanic
rocks (gray to brown in color) that contain very fine grained(<0.2 mm) acicular plagioclase. Devitrified scoria and small amountsof trachytic rocks containing abundant aligned acicular plagioclaseare also observed. Heavy minerals consist mainly of subhedral to
euhedral clinopyroxene and opaque minerals. Irregularly shaped
serpentinite grains, presumably altered from olivine, are also present.
Mon-AT-01 Upper horizon of the
Monshizu Formation,Nemuro Group onthe Konsen coastThe Ponporoto coast,
Hamanaka(43
/C148056.8400N,
145/C1410031.0700E;
Figure 6e)This sample was taken from a 3.0-m-thick vitreous
acidic tuff bed (Mon-AT) (Figure 7f). The fractured
surface of the sample is glassyThis wacke sandstone consists mostly of volcanic glass fragments and
minor amounts of quartz and plagioclase crystal fragments 0.1 –
0.2 mm in size, as well as columnar euhedral zircons and opaqueminerals. Trace amounts of apatite, biotite, clinopyroxene, and
amphibole are also observed. Volcanic glass fragments are composed
of very-fine-grained devitrified materials. Some quartz fragmentsshow undulatory extinction. Plagioclase commonly shows albite andCarlsbad twinning. Because this sample is very fine grained and wellsorted, we consider it to be a subaqueous secondary volcanic ash-fall
deposit
Ham-AT-01 Basal horizon of the
HamanakaFormation, NemuroGroup on theKonsen coastThe Mouraito coast
outcrop, Mouraito,Hamanaka(43
/C14909.8300N,
145/C1414019.0100E;
Figure 6f)This sample was obtained from a 6.3-m-thick vitreous
acidic tuff bed (Ham-AT) interbedded with deep-seaturbidite strata (Figure
7e). The fractured surface of
the sample is glassy and includes a cream-coloredcrackThis wacke sandstone consists mostly of volcanic glass fragments with
minor amounts of quartz, plagioclase, and potassium feldspar crystalfragments 0.1 –0.2 mm in size. Minor amounts of columnar euhedral
zircons and opaque minerals, and trace amounts of apatite andbiotite are also observed. Volcanic glass fragments are composed of
very-fine-grained devitrified materials. Plagioclase and potassium
feldspar show albite twinning and microcline structure, respectively.We consider this very-fine-grained and well-sorted sample to be asubaqueous secondary volcanic ash-fall deposit
KS-HMt-01 Upper horizon of the
Hamanaka
Formation, Nemuro
Group on theKonsen coastThe east coast outcrop
of Mouraito Port in
Mouraito, Hamanaka
(43
/C14909.8300N,
145/C1414019.0100E;
Figure 6f)This sample was collected from a 20-cm-thick turbidite
sandstone bed directly below the thick dolerite
sheet that separates the Hamanaka and Akkeshi
Formations (Figure 7d)This wacke sandstone contains many angular particles and has the
modal composition 1.0% quartz, 22.0% feldspar, 62.1% rock
fragments, 2.8% heavy minerals, and 12.1% matrix. Feldspars are
mostly subhedral plagioclase with albite twins. Most rock fragmentsare gray to brown basaltic volcanic rocks that contain very-fine-grained (<0.1 mm) acicular plagioclase. Minor amounts of nearlytransparent trachytic rocks contain abundant aligned acicular
plagioclase crystals, and devitrified scoria is also observed. Heavy
minerals are mainly subhedral clinopyroxene, rounded or irregularlyshaped serpentinite (olivine pseudomorphs), and opaque minerals.
(Continues)NANAYAMA ET AL. 15 of 32
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TABLE 1 (Continued)
Sample number Sampling horizon Location Field description Observation under the microscope
KS-AKb-01 Basal horizon of the
Akkeshi Formationon the Konsen coastThe west coast outcrop
of Mouraito Port inMouraito, HamanakaTown (43
/C14
8052.7100N, 145/C14
13045.0500E;
Figure 6f)This sample was taken from a slump-folded turbidite
sandstone 20 cm thick (Figure 7c). The fractured
surface is greenish-gray in colorThis wacke sandstone contains 13.5% quartz, 21.4% feldspar, 36.9% rock
fragments, 10.7% heavy minerals, and 17.5% matrix. Quartz occurs assingle crystals, and most do not show undulose extinction. Feldsparsare mostly plagioclase (many exhibiting short columnar automorphism)
with a small amount of potassium fel dspar. Rock fragments comprise
30.7% volcanic rocks (mostly hyalo-ophitic andesite, some pilotaxiticdacite, and small amounts of scori a and intersertal basalt), 4.6%
sedimentary rocks, 0.5% chert, and 1.1% schist. Heavy minerals aremainly pyroxene and amphibole, with lesser amounts of epidote,
apatite, zircon, garnet, and opaque minerals
HN-YS2-03 Ys2 member of the
Yusenkyo Formationin the Shiranuka hillsThe riverbed outcrop
of the HonbetsuRiver near the ForestRoad Bridge No. 11,Honbetsu
(43
/C1410026.2500N,
143/C1442058.4200E;
Figure 6c)This sample was taken from a medium-grained
sandstone bed interbedded with thick coarse-grained sandstone and pebbly sandstone beds(Figure
7h), presumed to have been deposited in a
deep-sea environment as a turbidite The fractured
surface of the sample is black in colorThis wacke sandstone contains many angular particles and has the
modal composition 0.4% quartz, 12.3% feldspar, 65.5% rockfragments, 4.0% heavy minerals, and 17.8% matrix. Feldspars aremostly plagioclase, and many exhibit short columnar automorphism.Rock fragments comprise 0.2% sedimentary rocks and 65.3%
volcanic rocks (mostly basalts and andesites, including pumice and
scoria). Heavy minerals are generally scarce and consist ofclinopyroxene and hornblende
KT-SRb-01 Lowest Unit, Saroma
Group in the KitamihillsThe riverbed outcrop
of the KutonnicoloRiver, Misato, Kitami
(43
/C1454010.2800N,
143/C1448023.0800E;
Figure 6b)This sample was obtained from a medium-grained
turbidite sandstone just above the basalconglomerate. The sampled unit is a volcanic
sandstone containing abundant andesite fragments.
The fractured surface of the sample is black in colorThis wacke sandstone contains many angular particles and has the modal
composition 1.0% quartz, 23.0% feld spar, 57.2% rock fragments, 1.8%
heavy minerals, and 17.0% matrix. Feldspars are mostly plagioclase,
and many exhibit short columnar automorphism. All rock fragments are
volcanic (mostly green to red basalt and andesite, including pumice andscoria). Heavy minerals are generally scarce and consist ofclinopyroxene, hornblende, and opaque minerals
NK-SRb-01 Lowest Unit, Saroma
Group in the Kitami
hillsThe riverbed outcrop
of the Nikura River,
Nikura, Saroma
(44
/C140025.0700N,
143/C1452058.2900E;
Figure 6a)This sample was taken from the medium-grained
sandstone overlying the basal conglomerate
(Figure 7i). Late Cretaceous Inoceramus fragments
have been reported in this horizon. The fracturedsurface of the sample is black in colorThis wacke sandstone contains many angular particles and has the
modal composition 1.0% quartz, 22.0% feldspar, 62.1% rock
fragments, 2.8% heavy minerals, and 12.1% matrix. Feldspars are
mostly plagioclase, and many exhibit short columnar automorphism.All rock fragments are volcanic rocks (mostly green to red basalt andandesite, including pumice and scoria). Heavy minerals are generallyscarce and consist of clinopyroxene, hornblende, and opaque
minerals
ST-PS54-01 Kawaruppu Formation,
Nemuro Group in theShiranuka hillsThe riverbed outcrop,
upstream SatonbetsuRiver inNupukibetsu,Onbetsu, Kushiro
City (43
/C146013.3100N,
143/C144809.4600E;
Figure 6d)This sample was collected from the top of a thick
sandstone bed (SC3) just below the acidic tuff bed inthe upper Kawaruppu Formation (Figure
7b), for
which Shibata ( 1984 ) reported a K –Ar age of 54.0
± 1.7 Ma (Figure 4). The fractured surface of the
sample is greenish-grayThis wacke sandstone contains many angular particles and has the
modal composition 9.9% quartz, 14.9% feldspar, 44.7% rockfragments, 1.9% heavy minerals, and 28.6% matrix. Quartz occurs assingle crystals, many of which do not show undulose extinction.Feldspars are mostly plagioclase, and many exhibit short columnar
automorphism; a small amount of potassium feldspar is also present.
Rock fragments comprise 15.3% sedimentary rocks, 25.4% volcanicrocks (mostly basalts and andesites), and 4.0% plutonic rocks. Heavyminerals are generally scarce and consist of clinopyroxene andhornblende associated with chromian spinel, zircon, and garnet16 of 32 NANAYAMA ET AL.
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Twenty-seven of the 30 grains showed concordant ages (Figure 9)
forming a single age cluster ranging from 79.97 ± 2.97 Ma to 66.26± 2.10 Ma (27 grains). We calculated the weighted average age
(YGC3 σ) of 68.19 ± 0.41 Ma (Figure
11).
6.4 |KS-HMt-01
We obtained 1000 zircon grains from 0.20 kg of sandstone. This sam-
ple yielded light brown, homogeneous, columnal, euhedral zircons and
rounded zircon grains, some purple in color. Of the 116 grains ana-lyzed, 116 had concordant ages (Figure
9) ranging from 2.7 Ga to
66.8 ± 1.7 Ma. We obtained at least 22 age clusters: 2.7 Ga (1 grain),
2.5–2.4 Ga (5 grains), 2.2 –1.7 Ga (17 grains), 1.3 –1.0 Ga (3 grain),
542–502 Ma (2 grains), 477 Ma (1 grain), 417 Ma (1 grain), 400 –
394 Ma (2 grains), 357 –352 Ma (2 grains), 334 –332 (2 grains), 305 –
298 Ma (2 grains), 280 Ma (1 grain), 268 –261 Ma (2 grains), 251 –
240 Ma (8 grains), 222 –217 Ma (2 grains), 207 Ma (1 grain), 196 –
156 Ma (16 grains), 153 –130 Ma (8 grains), 119 –109 Ma (5 grains),
88.7–83.5 Ma (3 grains), 79.5 –74.3 Ma (10 grains) and 70.4 –66.8 Ma
(6 grains). Of the youngest six grains (70.4 –66.8 Ma) and calculated
the weighted average age (YGC3 σ) to be 68.3 ± 0.9 Ma (Figure 11).
6.5 |KS-AKb-01
We obtained 500 zircon grains from 0.20 kg of sandstone. This sam-
ple yielded light brown, homogeneous, columnal, euhedral zircons,and rounded zircon grains, some purple in color. Of the 103 grains
analyzed, all were concordant ages ranging from 2.7 Ga to 64.5
± 1.2 Ma (Figure
10), and we recognized at least 30 age clusters:
2.7 Ga (2 grains), 2.6 Ga (2 grains), 2.5 Ga (5 grains), 2.1 Ga (1 grain),
1.9 Ga (4 grains), 1.8 Ga (4 grains), 1.7 Ga (2 grains), 917 Ma (1 grain),
491 Ma (1 grain), 477 –473 Ma (3 grains), 447 –428 Ma (3 grains),
385–375 Ma (2 grains), 361 –350 Ma (5 grains), 345 –336 Ma
(4 grains), 325 –317 Ma (4 grains), 284 –275 Ma (3 grains), 264 –
246 Ma (6 grains), 231 –224 Ma (2 grains), 213 Ma (2 grains), 187 –
193 Ma (3 grains), 184 –174 Ma (6 grains), 162 –159 Ma (2 grains),
157–155 Ma (3 grains), 149 –146 Ma (3 grains), 142 Ma (1 grain),
123–118 Ma (2 grains), 93.4 Ma (1 grain), 84.2 Ma (1 grain), 80.8 –
74.2 Ma (15 grains), and 73.6 –67.4 Ma (7 grains). From the youngest
seven grains (73.6 –67.4 Ma) and calculated the weighted average age
(YGC3 σ) to be 71.0 ± 0.8 Ma (Figure 11).
6.6 |HN-YS2-03
We did not obtain any zircons from the 0.20 kg of sandstone crushed.
6.7 |KT-SRb-01
We did not obtain any zircons from the 0.20 kg of sandstone crushed.TABLE 1 (Continued)
Sample number Sampling horizon Location Field description Observation under the microscope
ST-KWt-01 Kawaruppu Formation,
Nemuro Group in theShiranuka hillsThe riverbed outcrop,
upstream SatonbetsuRiver inNupukibetsu,
Onbetsu, Kushiro
City (43
/C145032.1100N,
143/C1448055.7700E;
Figure 6d)This sample was collected from the bioturbated silty
sandstone bed at the top horizon of the KawaruppuFormation and just below the basal erosive surfaceof the Urahoro Group (Figure
7a). The fractured
surface of the sample is greenish-gray in colorSilt-sized quartz and feldspar grains (mostly plagioclase) are scattered in
a clay-mineral matrix (90%). No heavy minerals other than zircon andchromian spinel are observedNANAYAMA ET AL. 17 of 32
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0.0100.0120.0140.016
0 . 00 . 10 . 20 . 30 . 4207Pb / 235U206Pb / 238U708090100
NK-SRb-01
(n = 32)
0.01100.01140.01180.01220.01260.01300.01340.01380.0142
0.055 0.065 0.075 0.085 0.095 0.105 0.1157478828690NK-SRb-01
(n = 25)206Pb / 238U
207Pb / 235U
0.00960.01000.01040.01080.01120.0116
0.045 0.055 0.065 0.075 0.085 0.09562646668707274 Ham-AT-01
(n = 27)206Pb / 238U
207Pb / 235U0.00960.01000.01040.01080.01120.01160.0120
0.056 0.060 0.064 0.068 0.072 0.076 0.080 0.084 0.088
207Pb/235U206Pb / 238U
64687276Mon-AT-01
(n = 30)0.00980.01020.01060.01100.01140.0118
0.055 0.065 0.075 0.085 0.095646668707274206Pb / 238U
207Pb / 235UKS-HMt-01
(n = 6)
0.00.10.20.30.40.50.6
0 2 4 6 8 1 01 21 460010001400180022002600
KS-HMt-01
(n = 116)
207Pb / 235U206Pb / 238U
FIGURE 9 Concordia diagrams of (left) all zircon U –Pb analyses and (right) the youngest age cluster per sample in (top) KS-HMt-01 and (middle)
NK-SRb-01. In the bottom panels, all zircons from (left) Ham-AT-01 and (right) Mon-AT-01 are shown. Error ellipses represent 2 σuncertainties18 of 32 NANAYAMA ET AL.
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0.00.10.20.30.40.5
02468
207Pb / 235U206Pb / 238U
2006001000140018002200ST-PS54-01
(n = 41)
0.00650.00750.00850.00950.01050.0115
0.00 0.02 0.04 0.06 0.08 0.10 0.12207Pb / 235U206Pb / 238U
444852566064687276ST-PS54-01
(n = 21)
0.00.20.40.6
048 1 2 1 660010001400180022002600206Pb / 238U
207Pb / 235UKS-AKb-01
(n = 103)
0.00980.01020.01060.01100.01140.01180.0122
0.06 0.07 0.08 0.096466687072747678(n = 7)KS-AKb-01206Pb / 238U
207Pb / 235U
0.00780.00820.00860.00900.0094
0.046 0.050 0.054 0.058 0.062 0.0665254565860ST-KWt-01
(n=28)
0.00.20.40.6
048 1 2 1 660010001400180022002600ST-KWt-01
(n = 119/120)206Pb / 238U
207Pb / 235U
206Pb / 238U
207Pb / 235U
FIGURE 10 Concordia diagrams of (left) all zircon U –Pb analyses and (right) the youngest age cluster per sample in (top) ST-KWt-01, (middle)
ST-PS54-01, and (bottom) KS-AKb-01. Error ellipses represent 2 σuncertaintiesNANAYAMA ET AL. 19 of 32
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0246810
44 49 54 59 64 69 74 79Number of grainsST-PS54-01
(n = 21)
YGC3σ:
61.3 ± 0.9 Ma
206Pb / 238U age (Ma)0246810121416
62 64 66 68 70 72 74 75Number of grainsHam-AT-01
(n = 27)
YGC3σ: 68.2 ± 0.4 Ma
206Pb / 238U age (Ma)
0246810121416
63 65 67 69 71 73 75 77Number of grainsMon-AT-01
(n = 30)
YGC3σ:
69.6 ± 0.3 Ma
206Pb / 238U age (Ma)
70 75 80 85 90 95Number of grains
12468101214NK-SRb-01
(n =25)
YGC3σ:
80.5 ± 0.5 Ma
206Pb / 238U age (Ma)ST-KWt-01
(n = 28)
YGC3σ:
54.8 ± 0.2 Ma
02468101214
50 52 54 56 58 60 62Number of grains
206Pb / 238U age (Ma)01234
64 66 68 70 72 74 76 78
206Pb / 238U age (Ma)Number of grainsKS-HMt-01
(n = 6)
YGC3σ:
68.3 ± 0.9 Ma
0123
64 66 68 70 72 74 76 78 80
206Pb / 238U age (Ma)Number of grainsKS-AKb-01
(n = 7)
YGC3σ:
71.0 ± 0.8
Ma
FIGURE 11 Histograms of the youngest zircon U –Pb age clusters in each sample; “n”indicates the number of grains in each cluster20 of 32 NANAYAMA ET AL.
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6.8 |NK-SRb-01
We obtained 60 zircon grains from 1.1 kg of sandstone. Although this
sample contained only a small number of zircons, they were uniquely
semi-automorphic to superficial debris grains with a homogeneous
brown color. Of the 30 grains analyzed, 26 were concordant, withages ranging from 90.4 ± 2.8 Ma to 76.9 ± 3.1 Ma (Figure
9). Exclud-
ing the oldest of these (grain 18, 90.4 Ma) as being beyond the 3 σ
error range, the remaining 25 grains formed a single cluster with ages
ranging from 84.1 to 76.9 Ma, and we calculated the weighted aver-age age (YGC3 σ) to be 80.5 ± 0.5 Ma (Figure
11).
6.9 |ST-PS54-01
We obtained 500 zircon grains from 0.3 kg of sandstone. This sample
yielded light brown, homogeneous, columnal, euhedral zircons, and
rounded zircon grains, some purple in color. Of the 35 grains analyzed,
34 were concordant (Figure 10). The ages ranged from 1.9 Ga
(1883 Ma) to 53.4 ± 5.5 Ma, comprising at least eight clusters: 1.9 Ga(1883 Ma, 1 grain), 338.6 –355.4 Ma (2 grains), 214.7 Ma (1 grain),
198.1 Ma (1 grain), 172.0 –181.6 Ma (2 grains), 85.6 Ma (1 grain),
66.2–79.1 Ma (7 grains), and 53.4 –65.7 Ma (21 grains). The weighted
average age (YGC3 σ) of the youngest 21 grains was calculated to be
60.9 ± 0.9 Ma (Figure
11).6.10 |ST-KWt-01
We obtained 10 000 zircon grains from 0.2 kg of sandstone. This sam-
ple yielded light brown, homogeneous, columnal, euhedral zircons, and
rounded zircon grains, some purple in color. Of the 120 grains analyzed,
28 were concordant (Figure 10). The ages ranged from 2.7 Ga to 52.9
± 1.3 Ma, comprising at least nine clusters: 2.7 Ga (1 grain), 2.0 Ga(2 grains), 733 Ma (1 grain), 645 Ma (1 grain), 498 –491 Ma (2 grains),
479 Ma (1 grain), 460 Ma (1 grain), 387 Ma (1 grain), 373 –368 Ma
(2 grains), 357 –352 Ma (2 grains), 346 –335 Ma (3 grains), 260 –253 Ma
(3 grains), 209 –2 0 5M a( 2g r a i n s ) ,1 9 5 –172 Ma (10 grains), 161 –153 Ma
(2 grains), 131 Ma (1 grain), 117 Ma (1 grain), 82.7 –7 6 . 7M a( 3g r a i n s ) ,
70.6–57.4 Ma (48 grains), and 57.1 –52.9 Ma (28 grains).
The weighted average (YGC3 σ) of the youngest 28 grains
(57.1 –52.9 Ma) was calculated to be 54.8 ± 0.2 Ma (Figure
11).
7|DISCUSSION
7.1 |Two reliable depositional ages in the lower
horizon of the Nemuro Group
The ages of the zircon U –Pb age clusters obtained from the vitric acidic
tuff beds in the upper horizon of the Monshizu Formation (Mon-AT) and
the lowest horizon of the Hamanaka Formation (Ham-AT) in the Nemuro
Early stage of the Kuril arc vocanism
Second collision, bend and uprift
First collision and upriftShiranuka hills Konsen coast
Fault
Fault
?Kiritappu Fm.Furubanya Unit
Konbumori Fm.
Shiomi Fm.
Senpohshi Fm.
Oborogawa Fm.
Monshizu Fm.
Ohtamura Fm.Hamanaka Fm.Tokotan Fm.50
60
70
80
Eocene PaleoceneCretaceous Paleogene
Campanian MaastrichtianKitami hills
Upper to
Lower units
Lowest
UnitSaroma Gr. Kushiro -
AkkeshiNemuro -
Hamanaka
?Fault
Fault
FaultKatsuhira
Fm.Kawaruppu
Fm.
Yusenkyo Fm.
(Saroma Gr.)Nemuro Gr.
Nemuro Gr. Nemuro Gr.
K/Pg
boundary
66 Ma40Urahoro Gr.
Rikubetsu Fm.
68-65 K-Ar40-39
81-80 (NK-SRb-01)55-55 (ST-KWt-01)
62-60 (ST-PS54-01)Urahoro Gr.
70-69 (Mon-AT-01)69-68 (Ham-AT-01)Akkeshi Fm.
69-67 (KS-HMt-01)
72-70 (KS-AKb-01)Ma
30
Futamata Andesite Oligocene
Nuibetsu Fm.
(Onbetsu Gr.)
Nokkamappu Fm.Prasolov magmatic complexes 30Danian Seland. Than. Ypresian Lutetian Bartn. Priabon.
?
??
Sphenoceramus
schmidti Zone
79-81?
RelativeProbability
206Pb / 238U age( M a )
1030501003005001000200010305010030050010002000
0.02.04.06.08.010.012.0
NK-SRb-01ST-KWt-01
ST-PS54-01Ham-AT-01
Mon-AT-010.02.04.06.08.010.012.0
54.9
68.269.5
KS-HMt-01KS-AKb-01
500100020000.02.0
1.000
ST-KWt-01
ST-PS54-01
KS-HMt-01KS-AKb-0180.4
68.367.9 80.461.461.4
67.9
68.3(b) (a)
FIGURE 12 (a) Geochronological correlation chart comparing sedimentary rocks of the Nemuro and Saroma groups in the Kitami hills,
Shiranuka hills, and along the Konsen coast. Modified from Shibata and Tanai ( 1982 ), Shibata ( 1984 ), Kaiho ( 1984 ), Okada et al. ( 1987 ), Nifuku
et al. ( 2009 ), Ueda ( 2016 ), De Grave et al. ( 2016 ), Nanayama et al. ( 2021 ), and Yabe et al. ( 2021 ). Green-colored formations do not contain
rounded zircons of NE Asian origin. Gr, group, Fm, formation. The red number indicates the reported zircon U –Pb ages (Ma) and the red number
associated with red stars and sample numbers indicate the new zircon U –Pb ages (Ma) obtained in this study. (b) Comparison of seven zircon U –
Pb age distributions measured in this studyNANAYAMA ET AL. 21 of 32
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Group on the Konsen coast are 70 –69 and 69 –68 Ma, respectively.
Moreover, we consider these ages to be highly reliable because the euhe-dral zircons show little wear, indicating their deposition shortly following
ash fall at both sites, and because the analytical ages are concentrated
around a single mode in both cases. Results of “ImageJ fiji ”analysis also
show that the zircons in these samples exhibit a wide range of aspect
ratios and essentially low roundness (Table S9, Figure S2, S3). Therefore,
these two depositional ages clarify the following three deductions.
1. Although the depositional age of the Monshizu Formation was previ-
ously estimated to be Maastrichtian (Matsumoto,
1970 )o re a r l y
Maastrichtian (Naruse et al., 2000 ), the upper depositional age of
70–69 Ma obtained for Mon-AT-01 indicates that the deposition of
the Monshizu Formation continued into the late Maastrichtian.
2. Because the K –Pg boundary is assumed to be between the upper
and middle members of the Akkeshi Formation, our lower limit on
the depositional age of the Hamanaka Formation (69 –68 Ma
obtained for Ham-AT-01) requires that the Hamanaka Formationcomprises exclusively upper Maastrichtian deposits. However,
Naruse et al. (
2000 ) reported Gaudryceras sp. in this formation andestimated that the depositional age was early Maastrichtian. We
conclude that this fossil age should be reinterpreted as reworked.In contrast, the reported K –Ar ages of the dolerite sheet that
intruded the Hamanaka Formation (65.5 ± 2.1 and 67.7 ± 2.1 Ma;
Shibata,
1984 ) are consistent with our zircon U –Pb age for Ham-
AT-01, indicating igneous activity during deposition.
3. The depositional age of the Oborogawa Formation (stratigra-
phically between the Monshizu and Hamanaka Formations)must therefore be between 70 and 68 Ma; hence, all strata of
this formation must also be upper M aastrichtian deposits. Shi-
bata's (
1986 )R b–Sr ages of ca. 70 Ma for the monzonite intrud-
ing the Oborogawa F ormation are thus consistent with our
zircon U –Pb ages.
7.2 |Stratigraphically unreliable zircon U –Pb ages
obtained for the Nemuro Group and their causes
The youngest age clusters obtained for the turbidite sandstones of the
upper horizon of the Hamanaka Formation (KS-HMt-01) and the lowest
0102030405060708090
0 50 100 150 200 250 300 350 400 450
U-Pb age (Ma) Older data:
2.6 Ga (N=1)
1.9 Ga (N=1)
1.8 Ga (N=1)NumberZones S, U, R
and Y (NHB)
0102030405060708090
0 50 100 150 200 250 300 350 400 450
U-Pb age (Ma)
Northern Unit
Southern Unit Older data:
1.9 Ga (N=1)
2.0 Ga (N=31)
2.7 Ga (N=1) 678 Ma (N=1)
833 Ma (N=1)
1.7 Ga (N=1) 1.8 Ga (N=1)
2.0 Ga (N=2) NumberNakanogawa
Group (SHB)
Iwaonai Flysch (Zone S)
Kamiokoppe Formation (Zone U)
Rurochi Formation (Zone R)
Kanayama Formation (Zone Y)
Nakazono Formation (Zone Y) 0102030405060708090
0 50 100 150 200 250 300 350 400 450
U-Pb age (Ma) Number
Ham-AT-01 Lower Nemuro Group
and Saroma Group
Mon-AT-01
NK-SRb-01
0102030405060708090
0 50 100 150 200 250 300 350 400 450U-Pb age (Ma) Older data:
460 Ma (N=1) , 479 Ma (N=1), 491 Ma (N=1), 498 Ma
(N=1), 645 Ma (N=1), 733 Ma (N=1), 2.0 Ga (N=2), 2.7 Ga (N=1)
1.9 Ga (N=1) 477-473 Ma (N=3), 491 Ma (N=2), 917 Ma (N=1),
1.7 Ga (N=2),1.8 Ga (N=4), 1.9 Ga (N=4), 2.1 Ga
(N=1), 2.5 Ga (N=5), 2.6 Ga (N=2), 2.7 Ga (N=2)
471 Ma (N=1), 505 Ma (N=1), 563 Ma (N=1), 1.3 Ga
(N=1), 1.8 Ga (N=3),1.9 Ga (N=8), 2.0 Ga (N=4),
2.1 Ga (N=1), 2.2 Ga (N=1), 2.4 Ga (N=1),
2.5 Ga (N=4), 2.7 Ga (N=1)Number
Upper Nemuro GroupST-KWt-01
ST-PS54-01
KS-AKb-01KS-HMt-01
FIGURE 13 Zircon U –Pb age distributions from all available data used to infer the provenance of (top left) the lower Nemuro group and
Saroma group; (top right) the upper Nemuro group; (bottom left) zones S, U, R, and Y in the northern Hidaka Belt (NHB); and (bottom right) theNakanogawa group in the southern Hidaka Belt (SHB; Nanayama et al.,
2018 ,2019 ). Modified from Nanayama et al. ( 2021 )22 of 32 NANAYAMA ET AL.
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a. Initiation subduction (90-85 Ma)
PKAb. Accreted Nikoro Group and covered by forearc basin sediments (85-80 Ma)Upwelling depleted mantleSinking slab
and rapid rollback
PKA OAC+NE Asia
?First collision and upliftc. First collision and uplift
(70 Ma)Ealy stage of the Nemuro forearc basin
NIkoro Group
(Greenstone complex) Saroma and
Nemuro Grs.Izanagi Plate
Izanagi Plate
Izanagi Plate Alkali basalt
Old ziricon grains85-80 Ma
(Campanian)
Olyutosky
Arc?PKA
Nikoro Group20 cm/yr
Izanagi Plate70 Ma
(Maastrichtian)
20 cm/yr
Izanagi
Plate
NIkoro Group
(Greenstone complex)
Tokoro metamorphic rock
ESAVBOCVB
Sino-Korean
CratonNorth Asian
Craton
PKAOAC
Amur microplateFirst collisionEurasia
PlateNorth Amerian Plate
FIGURE 14 Legend on next page.NANAYAMA ET AL. 23 of 32
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horizon of the Akkeshi Formation (KS-AKb-01) are markedly older than
those obtained for Mon-AT-01 and Ham-AT-01 (Figure 12).
Kaiho ( 1984 )r e p o r t e da ne a r l y –middle Eocene (early Lutetian, 47.8 –
41.3 Ma) Subbotina frontosa fossil in the uppermost horizon of the Kawar-
uppu Formation along the Satonbetsu River. However, the U –Pb age of
55.3 ± 0.4 Ma obtained herein for ST-KWt-01 (from the same horizon) is
about 6 –14 Myr older than the fossil age, which may reflect reworking of
older zircon grains from the lower horizon of the Nemuro Group. Similarly,the age obtained for sample ST-PS54-01 collected from bed SC3 just
below the acidic tuff bed in the upper Kawaruppu Formation is 60.9
± 0.9 Ma; this age must also be too old, by about 9.5 –4.3 Myr.
In these cases, we infer that it was not possible to obtain an accurate
depositional age because the sediments were diluted by recycled grains
during turbidite sedimentation. The number of zircon analyses was most
likely insufficient to accurately estimate the depositional age (Figure
12).
7.3 |When and where did the PKA form?
We obtained a zircon U –Pb age of 80.5 ± 0.5 Ma for sandstone sam-
ple NK-SRb-01 from the Lowest Unit of the Saroma Group. Reportedfragments of Inoceramus fossils in this unit (Sakakibara &
Tanaka,
1986 ) very smilar to Inoceramus schmidti , indicating deposi-
tion during the middle Campanian (81.0 –79.4 Ma; Shigeta &
Tsutsumi, 2018 ). These fossil ages are consistent with our results.
However, grain 18 from sample NK-SRb-01 returned an age of
90.4 Ma (Turonian); the reason for this outlier age is not clear, requir-
ing more age data for this sample.
We did not obtain any zircon grains from the Nokkamappu For-
mation. However, Okada et al. ( 1987 ) documented middle –late Cam-
panian nannofossils (80.5 –72.1 Ma) in the Nokkamappu Formation,
and the Sphenoceramus schmidti zone (81.0 –79.4 Ma, middle Campa-
nian) has been reported in this same outcrop (Matsumoto, 1970 ). In
contrast, Shibata ( 1986 ) reported slightly younger K –Ar ages of
76–72 Ma (late Campanian –early Maastrichtian) for intrusive rocks in
the Nokkamappu Formation at Cape Nosappu. In this case, K –Ar age
is a cooling age of ~350/C14C, so the intrusive age of the monzonite plu-
tonic rocks at Cape Nossapu should be older than that. Therefore, it isnot inconsistent with the fossil age.
We were unable to find zircon grains in the sandstones collected
from the Nokkamappu Formation (NM-NKK-01), the Yusenkyo For-mation (HN-YS2-03), and the Lowest Unit of the Saroma Group along
the Kunnikoro River (KT-SRb-01). Indeed, the zircon U –Pb age histo-
gram of the sandstones of the Nemuro and Saroma Groups reveals agap in magmatism prior to around 85 Ma (Figure
13). Therefore, we
consider that (1) there was no igneous body feeding sediments tothese sites before ca. 85 Ma, and (2) not enough time had passed for
zircon grains to be eroded out of the PKA igneous rocks. This isimportant evidence that, at that time, the PKA was an immature island
arc, such as an intra-oceanic arc on the Izanagi Plate, that could not
receive clastic sediment inputs. Furthermore, the ages of 70 –69 and
69–68 Ma for the vitric acidic tuff beds in the upper horizon of the
Monshizu Formation (Mon-AT-01) and the lowermost horizon of the
Hamanaka Formation (Ham-AT-01), respectively, were obtained fromconsistently columnal and euhedral zircons, as were those younger
than 85 Ma obtained from the Nemuro Group above the horizon of the
Hamanaka Formation (samples KS-HMt-01, KS-AKb-01, ST-PS54-01,and ST-KWt-01). It is reasonable to conclude that these fresh zirconswere derived directly from the PKA, suggesting that the PKA existed as
an intra-oceanic arc continuously until ~70 Ma (Figures
14and15). Our
estimation of the onset time of the PKA is also consistent with Vaeset al. (
2019 ) and Kimura et al. ( 2019 ).
Yamasaki and Nanayama ( 2018 ) suggested that the Daimaruyama
greenstones (Figure 3), large olistolith blocks in the southern part of
the Hidaka Belt, originated from an immature ocean arc. This claim
provides an important basis for our hypothesis that the subduction of
the PKA was initiated during 85 –80 Ma (see Section 7.4). In contrast,
the adjacent Tachiiwa greenstone body (Figure 3) was reported to be
an ocean-island basalt-like camptonite resulting from hotspot-like
igneous activity on the Izanagi Plate (Yamasaki & Nanayama, 2020 ).
Nonetheless, the amphibole K –Ar age for this body is ca. 95 Ma
(Owada et al., 1992 ), consistent with the initiation of subduction dur-
ing 85 –80 Ma. Similarly, the greenstone complex of the Nikoro Group,
previously described as fragments of the Tokoro island arc(e.g., Ueda & Miyashita,
2005 ; Zharov, 2005 ), more likely originated
from the subduction of seamounts or an oceanic plateau (Sakai
et al., 2019 ; Sakakibara et al., 1986 ; Yamasaki & Nanayama, 2017 ),
based on the existence of the high-pressure metamorphic rocks that
were generated contemporaneously along the eastern edge of the
Tokoro Belt (Figure 3).
7.4 |Timing of the PKA collision with NE Asia and
subsequent tectonics
Here, we discuss the origins of the detrital zircons older than 85 Ma
in the Hamanaka, Akkeshi, and Kawaruppu Formations (samples KS-HMt-01, KS-AKb-01, ST-PS54-01, and ST-KWt-01). Kiminami (
1979 )
investigated the sandstone composition of the Nemuro Group along
the Konsen coast based on thin-section observations and reportedthat detrital zircon grains appear from younger formations above the
Hamanaka Formation. However, he did not describe in detail whether
FIGURE 14 Schematic tectonic model of NE Asia, Hokkaido, and Sakhalin. Upper panel: Paleogeographic maps of NE Asia, Hokkaido, and
Sakhalin during the late cretaceous (85 –70 Ma). Large black arrows indicate oceanic plate motions (modified from Vaes et al., 2019 ). OCVB,
Okhotsk –Chukotka volcanic belt; ESAVB, east Sikhote –Alin volcanic belt; OAC, Okhotsk arc complex. (a) The initiation of subduction during 90 –
85 Ma. (b) The covering of the accreted oceanic plateau by forearc basin sediments during 85 –80 Ma. (c) The first collision and uplift stage
around 70 Ma. PKA, Paleo –Kuril arc24 of 32 NANAYAMA ET AL.
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?Bend and uplift
20 cm/yr
Izanagi PlateIPRIdPaleo-Kuril ArcEast Sikhote-Alin Arc
Collision Complex
of the Hidaka SupergroupUrahoro
Group(40-39 Ma)55–46 Ma
(Early Eocene)40-39 Ma
(Late Eocene)
7 cm/yr
Pacific PlateYubetsu and
Nakanogawa
Groups
In situ basaltSecond
collision
7 cm/yr
Pacific PlateTk TkNmKm
YzEurasia
PlateEurasia
PlateNorth Amerian
Plate
PKAOAC+NE Asia
?d. Deposition of deep-sea fan complex (55-46 Ma)
Izanagi Plate ˠOld detrital ziricon grains
NIkoro
GroupNakanogawa and
Yubetsu Grs.
e. Second collision, bent and uplift (40 Ma)
NIkoro
Group OAC+NE AsiaNakanogawa and
Yubetsu Grs.
Nemuro Group PKA
?Urahoro
Group
Ultramafic rockDetrital chromian spinelsYubetsu and
Nakanogawa Groups
Second collision, bend and upliftOld detrital ziricon grainsSubduucted IPR
FIGURE 15 Legend on next page.NANAYAMA ET AL. 25 of 32
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these zircon grains were euhedral or rounded. Nakazoe ( 1963 ) sys-
tematically examined the heavy-mineral compositions of sandstonesof the Nemuro Group in the Shiranuka hills and all of the Katsuhira
Formation and the Kawaruppu Formation (upper Maastrichtian –
middle Eocene). He reported that all strata contained rounded purplezircon grains, and that sandstones in the Akkeshi and Tokotan Forma-tions along the Konsen coast also contain similarly rounded zircon
grains.
In this study, we did not find rounded zircon grains in the Lowest
Unit of the Saroma Group (samples KT-SRb-01, NK-SRb-01), the
Yusenkyo Formation (HN-YS2-03), the Nokkamappu Formation (NM-
NKK-01), the upper horizon of the Monshizu Formation (Mon-AT-01),or the lowest horizon of the Hamanaka Formation (Ham-AT-01). In
contrast, we found rounded zircon grains, some purple in color, in for-
mations above the Hamanaka Formation. All these zircons returnedU–Pb ages of 1 –3 Ga. The image analysis show that these detrital zir-
con grains clearly rounded compared to younger (70 –68 Ma) euhedral
zircons, and supports microscopic observations (Figure S2d,e;
Table S9).
These older grains can be reasonably interpreted by considering
depositional histories. After the PKA had been established, it existed
as an oceanic island arc until ~70 Ma, when the lowermost horizon ofthe Hamanaka Formation (Ham-AT) was deposited. Since that time,
the PKA must have collided with NE Asia and transitioned to a conti-
nental arc because older detrital zircon grains could only have beenderived from cratonic crust. Reasonable candidates are the North
Asian Craton, the Kolyma –Omolon rock body (Archean –Jurassic,4G a –145.0 Ma) in the Chukotka region along the northern Sea of
Okhotsk, and the Okhotsk cratonal terrane (Archean –Proterozoic,
4–0.5 Ga) in the Magadan region (Figure
1). Further research is
required to distinguish the provenance of these recycled zircons.
Vaes et al. ( 2019 ) estimated the PKA and the Olyutorsky arc had
collided with NE Asia around 55 –45 Ma according to geological infor-
mation around Kamchatka Peninsula and Sakhalin Island (Hourigan
et al., 2009 ; Kostantinovskaya, 2001; Parfenov et al., 2011 ). Kimura
et al. ( 2019 ) also estimated the PKA had collided with NE Asia around
ca. 50 –40 Ma on the plate motion model of Seton et al. (2015). As far
as we know, there is no geological evidence of a large-scale collision
at 55 –45 Ma in eastern Hokkaido. On the other hand, their hypothe-
sis contradicts the idea that the Izanagi Pacific Ridge passed at
55–47 Ma during the deposition in the Hidaka Belt in Hokkaido
(Nanayama et al., 2021 ). We guess the PKA had collided with NE Asia
around 70 Ma, and we consider that the area including the Sea ofOkhotsk and Kamchatka separated from the Eurasian Plate to form
the Okhotsk microplate and began to behave as part of the North
American Plate after the first collision. The PKA was then sutured tothe Eurasian Plate with a dextral strike-slip component, forming the
Hidaka Belt (Nanayama et al.,
2021 ). It is thought that this event
established the underlying structure of central –eastern Hokkaido.
The large-scale slump beds developed in the Akkeshi Formation
and the thick, coarse conglomerate bed containing boulder clasts
developed in the Kiritappu Formation indicate a change in deposi-tional environment caused by the full-scale collision of the PKA and
NE Asia, and we here reinterpret this change as evidence for the uplift
FIGURE 15 Schematic tectonic model of NE Asia, Hokkaido, and Sakhalin. Upper Paleogeographic maps of NE Asia, Hokkaido, and Sakhalin
during the early and late Eocene (55 –46 and 40 –39 Ma, respectively; modified from Nanayama et al., 2021 ). Large black arrows indicate oceanic
plate motions (modified from Vaes et al., 2019 ). Km, Kamuikotan metamorphic zone; Tk, Tokoro Belt; Nm, Nemuro Belt; Id, Idonnappu zone; Yz,
Yezo forearc basin; IPR, Izanagi –Pacific ridge. (d) Deposition of the deep-sea fan complex stage during 55 –46 Ma. (g) The second collision of the
PKA and NE Asia and associated bending and uplift during 40 –39 Ma. PKA, Paleo-Kuril arc; OAC, Okhotsk arc complex
Yusenkyo Formation (Saroma Group)Onbetsu Group
Urahoro Group
Kawaruppu Formation
(Nemuro Group)
Senpohshi Formation
(Nemuro Group)Modern Kushiro City
Akkeshi Formation
(Nemuro Group)
1015202530354045505560
0 1 02 03 04 05 06 07 08 0Latitude (°)
Age (Ma)FIGURE 16 Paleolatitudes of sedimentary
rocks deposited after the Yusenkyo formation(Saroma group) calculated from data reported byFujiwara and Kanamatsu (
1990 ), Fujiwara et al.
(1995 ), Kanamatsu et al. ( 1992 ), Katagiri et al.
(2019 ), and Nifuku et al. ( 2009 )26 of 32 NANAYAMA ET AL.
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of the Nemuro Peninsula. However, in the Shiranuka hills, such crustal
movements and uplift processes are represented by upward coarsen-ing from the Katsuhira Formation, mainly hemipelagic mudstone, to
the Kawaruppu Formation, mainly turbidite sandstone. This change in
depositional environment occurred after the Selandian (61.6 –
59.2 Ma), slightly after uplift along the Konsen coast (Kaiho,
1984 ).
The ages of detrital monazites (Yokoyama, 2016 ) and zircons
(Murakami et al., 2016 ) reported for the Lower and Middle Units of
the Saroma Group (Maastrichtian; Murakami et al., 2016 ) suggest that
they, like the formations above the Hamanaka Formation in the
Nemuro Group, may derive not only from the PKA, but also from amuch older geological body. Indeed, we extracted columnal euhedralzircons dating to 85 –75 Ma from the Lowest Unit of the Saroma
Group. This result indicates that changes in the sediment sources of
the Saroma and Nemuro Groups are linked with each other, i.e., thatthese two groups formed the same sedimentary basin (forearc or
intra-arc) of the PKA.
7.5 |What was the Okhotsk paleoland?
The geology of the Sea of Okhotsk, and our understanding of it, dif-
fers greatly between its northern and southern parts. Little geological
information is available for the northern part of the Sea of Okhotsk;most seismic exploration data were obtained in the 1980s and are ofpoor resolution. It has been argued that the crust in this area has
thinned to approximately 20 km thick (Gnibidenko &
Khvedchuk,
1982 ; Xu et al., 2016 ; Figure 1); however, no deep-sea
drilling data of the basement have been obtained to confirm this,
which should be addressed by future seismic reflection imaging and
dredge sampling of the seafloor basement. Gnibidenko and Khved-chuk (
1982 ) reported K –Ar ages of granites and granodiorites in
dredge samples from the Okhotsk Basin that broadly span 209 –
68 Ma. These ages are too young for the basement rocks of the Seaof Okhotsk to be a microcontinent; these rocks must instead comprisemultiple Late Triassic –Late Cretaceous volcanic arcs (Okhotsk arc
complex; Figures
14, 15 ). Indeed, the Kvakhona arc terrane and the
Sredinny massif exposed on the west coast of the Kamchatka Penin-sula have consistent Middle Jurassic –Early Cretaceous
(Konstantinovskaya,
2001 ) and Late Cretaceous ages (Hourigan
et al., 2009 ), respectively, and their western extensions continue to
the seafloor basement of the Sea of Okhotsk (Xu et al., 2016 ).
In contrast, the southern part of the Sea of Okhotsk is charac-
terized by the Kuril arc and backarc basin. Xu et al. ( 2016 )
described the Sea of Okhotsk area as widely subaerial and covered
by fluvial deposits with coal seams during the Paleocene to Eocene.
Once the modern Kuril arc had been established in the lateEocene –early Oligocene (De Grave et al.,
2016 ;V a e se ta l . , 2019 ;
Figure 2), the area gradually subsided during the Oligocene –early
Miocene to become the present-day Sea of Okhotsk. The Kuril
backarc basin then opened during a main episode of extension dur-ing the late Oligocene to early –middle Miocene (Kimura &Tamaki,
1985 ; Nanayama, Kurita, et al., 2020 ; Nanayama,
Watanabe, et al., 2020 ), which had ended by the late Miocene (9 –
7M a ;I k e d ae ta l . , 2000 ;X ue ta l . , 2016 ).
The Academy of Sciences Plateau is a large bedrock rise to the
north of the Kuril backarc basin (Figure 2). Various rock types have
been reported in dredge samples co llected there. In particular,
g r a n i t ea n da n d e s i t es a m p l e sh a v eL a t eC r e t a c e o u sK –Ar ages of
95–75 Ma (Burk & Gnibidenko, 1977 ). Some of these samples may
be ice-rafted debris, but most are considered to have been in situ.
Geodekyan et al. ( 1977 ), Burk and Gnibidenko ( 1977 ), and Kimura
and Tamaki ( 1985 ) noted that igneous rocks exposed near the
Academy of Sciences Plateau are the same as those intruding andextruding the Nemuro Group along the Konsen coast. It is there-
fore likely that the expansion of the Kuril backarc basin divided the
remnant of the PKA into northern and southern parts that nowoccur in parallel orientations to the north and south of the basin,
respectively (Figure
2), as is the case for the modern Izu –Bonin –
Mariana arc and Kyushu –Palau Ridge (remnant arc) on the Philip-
pine Sea Plate (Nishizawa et al., 2016 ).
7.6 |New tectonic model for the evolution of the
Kuril arc region
Based on literature data (Fujiwara et al., 1995 ; Fujiwara &
Kanamatsu, 1990 ; Kanamatsu et al., 1992 ; Katagiri et al., 2019 ; Nifuku
et al., 2009 ), we calculated the paleolatitudes of the Kuril arc region
(Figure 16). The upper and lower confidence limits on the paleolati-
tude were calculated using formulas described in Butler ( 1992 ). The
mean calculated paleolatitudes show a gradual increasing, which sug-
gest that the Kuril arc region has moved to higher latitudes since thedeposition of the Upper Cretaceous Yusenkyo Formation. However,
confidence limits of averaged paleolatitudes are widely range, the
minimum difference of paleolatitude between the oldest data andyoungest data is calculated to confirm a reasonability of our interpre-tation. The upper limit of paleolatitude of Yusenkyo Formation
(70–75 Ma) is 35.4
/C14. The lower limit of paleolatitude of Onbetsu
Group (32 –36 Ma) is 38.4/C14. These values lead that 3/C14increasing in lati-
tude from 70 to 32 Ma as minimum estimation. Because that is signifi-
cant long distance, we consider that the dislocation of Kuril arc region
to higher latitudes is a persuading interpretation.
Domeier et al. ( 2017 ) and Vaes et al. ( 2019 ) considered that the
PKA and the Olyutorsky arc were an oceanic island arc and occurred
on the northern edge of the Pacific plate. They assumed a subductionzone facing southward along the north side of these paleoarcs. Fur-
thermore, they considered these oceanic island arcs have been thrust
up to the Kamchatka Peninsula and Sakhalin Island around 55 –45 Ma
according to the geological information of Russian researchers(Hourigan et al.,
2009 ; Parfenov et al., 2011 ; Zhao et al., 2019 ). How-
ever, the greenstone accretionary complex (Nikoro Group) of the
Tokoro Belt in Hokkaido was misidentified, so they did not explainthe accretion to the PKA around 85 –80 Ma. We should assume aNANAYAMA ET AL. 27 of 32
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northward subduction zone along the south side of the PKA
(Nanayama et al., 1993 ,2021 ).
Incorporating these data into the discussion above, we propose
the following tectonic model for the evolution of the Kuril arc region
(Figures 14and15).
1. Around 85 Ma, a rupture occurred, and an initial subduction
began on the oceanic Izanagi Plate, establishing the earliestPKA as an oceanic island arc, most likely at lower latitude than
that of the modern Kuril arc (Kanamatsu et al.,
1992 ;V a e s
et al., 2019 ).
2. Around 85 –80 Ma, the greenstone body of the Nikoro Group was
accreted to the PKA. The Nokkamappu, Otamura, and Monshizu
Formations of the Nemuro Group were deposited on the forearc
side, and the Lowest Unit of the Saroma Group and the YusenkyoFormation were deposited above the Nikoro greenstone complex
by 81 –80 Ma. This accretion event also occurred at lower latitude
(Kanamatsu et al.,
1992 ; Vaes et al., 2019 ).
3. The PKA continued to move northward and began to collide with
NE Asia around 70 Ma. The Hamanaka to Akkeshi Formations of
the Nemuro Group were deposited on the forearc side, theLower to Upper Units of the Saroma Group were deposited
above the Nikoro greenstone c omplex, and large amounts of
clastic sediments derived from t he NE Asian continent were sup-
plied to the area. The collision was in full force by around 65 Ma,when large slump beds formed during the deposition of the
Akkeshi Formation. The Nemuro Group along the Konsen coast
experienced large-scale uplift, supplying abundant boulder claststo the Kiritappu Formation. The supply of these coarse materials
continued until ~50 Ma.
4. Around 55 –47 Ma, a large-scale submarine fan formed at the foot
of the collision between NE Asia and the PKA; this fan became the
Hidaka Supergroup that constitutes the Hidaka Belt.
5. Around 40 Ma, a second collisio n occurred, suturing the Nemuro
and Tokoro Belts together and bending them clockwise. TheNikoro Group experienced large-scale uplift and supplied boul-
ders and chromianspinel grains to the Urahoro Group (Nanayama
et al.,
1994 ).
6. The modern Kuril arc was established around 36 Ma (late Eocene –
early Oligocene). The Kuril backarc basin began expanding around
25 Ma, separated from the NE Asian continent by a large-scaleright-lateral displacement, and the Sea of Okhotsk region including
the PKA began to behave as an individual minor plate. After
10 Ma, the Kuril forearc sliver moved westward, initiating a thirdcollision (Kimura,
1986 ).
7. Finally, large-scale tectonic erosion in the Kuril Trench fragmen-
ted the PKA. Because both the eastern and northern(or western) edges are assumed to occur in Russia, this pointremains speculative. Little major progress has been made
toward understanding the foundation of the Sea of Okhotsk.
Future research on the coastal areas and seafloor of the Sea ofOkhotsk will be bolstered by cooperation between Japanese
and Russian researchers.8|CONCLUSIONS
The Nemuro and Saroma Groups distributed in eastern Hokkaido rep-
resent sediment deposited in the forearc basin of the PKA during the
Late Cretaceous to middle Eocene. We determined the U –Pb ages of
detrital zircons within these sandstones and interbedded tuffs and
drew the following conclusions.
1. The PKA likely originated as an oceanic island arc on the oceanic
Izanagi Plate around 85 Ma. The Nikoro greenstone complex was
accreted to the PKA just before 81 –80 Ma, and the Lowest Unit of
the Saroma Group was deposited on its upper surface.
2. The PKA subsequently collided with NE Asia around 70 Ma, when
the Hamanaka Formation was deposited, at which time the PKA
transitioned to a continental arc. This collision caused large-scaleuplift of the paleoarc, generating large slump beds during the
deposition of the Akkeshi Formation and providing large clasts to
the coarse-grained, thick conglomerate bed of the KiritappuFormation.
3. Although we do not have data directly bearing on why the North
American Plate was established in the edge of NE Asia, we specu-late that it separated from the Eurasian continent already ~70 Ma
when NE Asia first collided with the PKA. Subsequently, the PKA
has behaved as part of the North American Plate. ThePaleo-Japanese arc (East Sikhote –Alin arc) and the PKA subse-
quently joined along the Hidaka Belt around 40 Ma, when the Ura-
horo Group was deposited and bent clockwise due to the ongoing
collision.
4. The Kuril –Kamchatka arc basin was established around 36 Ma, and
the Kuril backarc basin opened in the late Oligocene to early Mio-
cene, resulting in the present-day tectonic setting.
ACKNOWLEDGMENTS
This paper is dedicated to Hakuyu Okada (Kyushu University),who was a pioneer in sandstone compositional studies in Hok-kaido. FN thanks Yoshiki Fujiwara (Hokkaido University), and
Kazuo Kiminami (Yamaguchi University) for raising the issue of
the PKA. We thank Jinichiro Maeda (Hokkaido Research Centerfor Geology), Seiichi Toshimitsu , Hirokuni Oda (Geological Survey
of Japan, AIST), and Shigeto Inokuma (Nemuro City Museum of
History and Nature) for their help with the details of Hokkaidogeology. Daniel Pastor-Galán, Gaku Kimura, Toru Takeshita and
an anonymous reviewer are greatly appreciated. This work was
supported by JSPS KAKENHI Grant Numbers JP19K04025 andJP21H04523.
CONFLICT OF INTEREST
The authors declare no conflict of interest.
ORCID
Futoshi Nanayama
https://orcid.org/0000-0001-8693-4784
Toru Yamasaki https://orcid.org/0000-0002-9300-0338
Toshiya Kanamatsu https://orcid.org/0000-0002-7108-453428 of 32 NANAYAMA ET AL.
14401738, 2022, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/iar.12458 by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Hideki Iwano https://orcid.org/0000-0001-5687-2105
Tohru Danhara https://orcid.org/0000-0002-0017-4161
Takafumi Hirata https://orcid.org/0000-0003-4683-9103
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ARTICLES
Deep structure of the northeastern Japan arc
and its relationship to seismic and
volcanic activity
Akira Hasegawa, Dapeng Zhao*, Shuichiro Hori, Akira Yamamoto
& Shigeki Horiuchi
Observation Center for Prediction of Earthquakes and Volcanic Eruptions, Faculty of Science, Tohoku University, Sendai 980, Japan
Evidence for possible deep-seated magmatic
activity beneath northeastern Japan can be
obtained from seismic observations. Tomographic
inversions of P-wave velocity data show low
velocity zones distributed in the crust and upper
mantle beneath active volcanoes. In or around
these zones, anomalously low-frequency micro
earthquakes, perhaps caused by magmatic activity,
are found at depths of 25-40 km; distinct S-wave
reflectors, corresponding to the upper surfaces of
magma bodies, are found at depths of 10-18 km.
Large crustal earthquakes also occur around the
low-velocity zones. These observations reveal the
distribution of magma reservoirs at depths and
elucidate its relation with shallow seismicity
beneath volcanic arcs.
* Present address: Geophysical Institute, University of Alaska Fairbanks,
Fairbanks 99775, USA
NATURE · VOL 352 · 22 AUGUST 1991 NoRTHEASTERN Japan is located at one of the most typical
subduction zones in the world, where the oceanic Pacific plate
plunges downward into the mantle under the continental plate.
Seismic activity is extremely high along the main thrust zone
beneath the landward slope of the Japan trench. This shallow
seismicity (shallower than 50-60 km) is characterized by low
angle thrust faulting1, which reflects directly the relative move
ments of the two converging plates and the downward motion
of the subducting plate. The shallow seismicity along the main
thrust zone merges continuously with the deep seismic zone
which dips beneath northeastern Japan. The deep seismic zone,
originating not at the boundary between the two converging
plates but within the subducted Pacific plate, is composed of
two thin planes, parallel to each other and 30-40 km apart2• The
upper seismic plane is characterized by down-dip compression,
whereas the lower plane by down-dip extension2• Formation of
the double seismic zone can be explained at least qualitatively
by the unbending of subducted oceanic plate3.
In addition to these seismicities, shallow earthquakes confined
mainly to the upper crust4 occur beneath the inland area.
Although the activity of these crustal events is less severe than
that along the main thrust zone, historically severe damage has
been inflicted by many crustal earthquakes because they are
close to populated areas. Many active volcanoes are distributed
in the land area, mainly along the volcanic front which runs
683
ARTICLES
through the middle of the land area nearly parallel to the
trench axis. Eruptions of these active volcanoes, although not
frequent, have been another source of natural disaster in this
country.
Here, we present a precise three-dimensional seismic velocity
structure of the crust and upper mantle beneath this volcanic
arc obtained by using data acquired through microearthquake
observations. Tomographic images of the estimated three
dimensional P-wave velocity structure show low-velocity (low
V) zones, which are continuously distributed from the upper
mantle to the upper crust beneath active volcanoes. Excep
tionally deep microearthquakes (25-40 km) of anomalously low
frequency (LF) are detected at nine locations around the low-V
zones beneath active volcanoes, suggesting a close relation with
the magmatic activity in this depth range. At shallower depths
(10-18 km) in or around the low-V zones, we find distinct
reflectors of S-waves. The locations of reflection points and the
calculated reflection coefficients suggest that the reflectors depict
the upper surfaces of magma bodies in the mid-crust.
P-wave velocity structure
Tomographic studies beneath the Japan islands have been made
by several researchers. Hirahara et al.5 investigated the P-wave
velocity structure beneath the central part of Japan, finding
low-V zones in the wedge portion of the upper mantle above
the high-V Philippine Sea and Pacific plates subducting beneath
central Japan. Hasemi et al.6 and Obara et al.7 investigated the
structure beneath the central part of northeastern Japan, and
found low-V zones in the crust and upper mantle beneath active
volcanoes. These results strongly suggest that the low-V zones
are closely related to arc volcanism and indicate the importance
of tomographic studies for understanding arc volcanism,
although they involve the problems described below.
Seismic velocity discontinuities (SVDs), such as the Moho
and Conrad discontinuities and the upper boundary of the
subducted oceanic plate, are known to exist in the crust and
upper mantle beneath the northeastern Japan arc and to have
large undulations8-11• Most tomographic studies made so far,
however, do not consider the effect of the complicated shapes
of the SVDs, or even their existence. Thus estimated tomographic
images include not only velocity variations but also the effect
of SVDs, especially in their vicinity. Another problem involved
in most of the conventional tomographic studies is that three
dimensional ray tracing is not used to calculate seismic ray paths
and travel times. In very heterogeneous regions such as subduc
tion zones, ray paths calculated from a simple one-dimensional
velocity model deviate considerably from the real paths for
seismic waves with large hypocentral distances; this seriously
distorts the final tomographic images.
A new method of seismic tomography has been developed to
solve these problems12
• This method is applicable to a general
velocity structure with many SVDs of complicated shapes in the
modelling space and with three-dimensional variations in the
seismic velocity in each layer bounded by these discontinuities.
Three-dimensional grid nets are arranged in every layer
individually, and the velocity at each grid point is taken to be
an unknown parameter. A velocity at any point in the modelling
space is calculated by linearly interpolating the velocities at
eight surrounding grid points. Velocity parameter derivatives
are calculated by using a linear interpolation function expressing
the velocity field. An efficient three-dimensional ray-tracing
algorithm has been developed which iteratively uses the pseudo
bending technique13 and Snell's law. This algorithm calculates
ray paths and travel times rapidly and accurately for seismic
waves in a complicated velocity structure. The algorithm
developed here can include in the inversion the arrival-time data
of later phases, such as reflected or converted waves at the SVDs,
which greatly improve the ray coverage in modelling space. A
conjugate gradient solver, the LSQR algorithm14, is used to solve
the extremely large and sparse system of observation equations
684 ansmg from the inversion problem. Hypocentre and velocity
parameters are determined simultaneously. Because of the
LSQR algorithm, it is unnecessary either to set up the normal
equation or to use the parameter separation technique, which
are very time-consuming and formidable when the number of
unknown parameters is large.
High-resolution P-wave tomographic images of the crust and
upper mantle beneath the Japan islands are obtained by applying
the method to arrival-time data observed by the seismic networks
of several national universities in Japan. In addition to first P
and S-wave arrivals, later arrivals of SP waves converted at the
Moho discontinuity, and PS and SP waves converted at the
upper boundary of the subducted Pacific plate, are read from
seismograms observed by the seismic network of Tohoku Uni
versity and are included in the inversion. We selected 512 shallow
events and 688 intermediate-depth and deep events which are
distributed uniformly in the upper crust and in the depth range
of 40-650 km along the Wadati-Benioff zones corresponding to
the subducted Philippine Sea and Pacific plates. In total, 50,919
data points corresponding to arrival times observed at 208 seis
mic stations distributed in the Japan Islands are collected from
the 1,200 events.
The medium under study is divided into four layers by the
Conrad and Moho discontinuities and the upper boundary of
the subducted Pacific plate, these four layers corresponding to
the upper crust, the lower crust, the wedge mantle and the mantle
beneath the upper boundary of the Pacific plate. The depth
distributions of these SVDs are fixed and expressed by con
tinuous functions of spatial locations so as to coincide with the
results of previous studies10•15'16• We arrange 1, 1, 16 and 16
layers of grid nets in the upper crust, the lower crust, the wedge
mantle and the mantle beneath the upper plate boundary, respec
tively. The separation between grid points is 25-30 km in both
vertical and horizontal directions. Three iterations are carried
out in the nonlinear inversion, as the difference in r.m.s. residual
of the inverted solutions between the second and third iterations
is sufficiently small. In total, a 48% variance reduction is
achieved. Checkerboard resolution tests17
'18 with various wave
lengths of velocity change are made to evaluate the resolution
of the tomographic images obtained. The original patterns of
the checkerboards are well reconstructed for the study area.
Reconstructed amplitudes of velocity anomalies are more than
80% of the original amplitudes for most grids shallower than
-200 km, and 60-70% for those deeper than that. The uncer
tainty of the inverted velocity perturbations is estimated to be
-0.5% for grids shallower than -200 km and less than 1% for
deeper grids. These indicate that a meaningful solution is accur
ately obtained.
We have updated the tomographic images of northeastern
Japan by Hasemi et al.6 and Obara et aC by covering a wider
area and by using higher resolution. Figure 1 shows the estimated
fractional P-wave velocity perturbations at a depth of 40 km
beneath northeastern Japan. As seen from this figure, most active
volcanoes are located in the low-V zones. Figure 2a shows an
east-west vertical cross-section. The low-V zone located just
beneath active volcanoes in the crust and the uppermost mantle
dips to the west in the wedge mantle and extends to a depth of
-150 km. This feature is common to all the low-V zones located
in the crust and the upper mantle beneath northeastern Japan.
Although in the cross-section shown in Fig. 2a, the low-V zone
seems to be split into two parts, one shallower (-50 km), the
other deeper ( -100 km), the other low-V zones beneath active
volcanoes continuously extend to a depth of -150 km; one
example is shown in Fig. 2b. A high-V zone corresponding to
the subducted Pacific plate is clearly seen. The thickness of the
high-V zone is estimated to be 80-90 km. This agrees with the
recent study by Umino et al.19 who detected a later phase
reflected from the bottom of the high-V Pacific plate beneath
northeastern Japan and estimated its thickness to be -85 km
from arrival-time analysis. Figure 2a also shows that the double
NATURE · VOL 352 · 22 AUGUST 1991
seismic zone is located in the upper half of the high-V Pacific
plate.
Low-frequency microearthquakes
More than 98% of shallow microearthquakes in the land area
of northeastern Japan occur in the upper 15 km of the crust.
The cut-off depth for these shallow events can be interpreted
as a brittle-ductile transition or a stick-slip to stable-sliding
transition due to increasing temperature with depth20-22• The
upper 15 km of the crust forms a brittle seismographic zone. A
systematic investigation of shallow inland microearthquakes
reveals however, that exceptionally deep events do occur
(although they are rare) at nine locations in northeastern Japan
in the depth range 25-40 km, which is well below the base of
the brittle seismogenic zone. The characteristics of these
anomalous events are as follows: their focal depths are
anomalously deep (25-40 km); the predominant frequencies of
both P-and S-waves are extremely low (1.5-3.5 Hz); the
maximum magnitude is at most 2.5; and they occur around
active volcanoes or around the low-V zones.
Examples of seismograms of these anomalously deep events
are shown in Fig. 3a. Figure 3b represents an event whose
epicentre is close to the event shown in a, but which has a
normal focal depth (10 km). What is obvious on comparing the
two seismograms is that this anomalously deep event has
extremely low predominant frequencies both for P-and S-waves;
this was found to apply to all the events at depths 25-40 km
found at the nine locations. Intermediate-depth earthquakes at
depths 70-150 km, occurring in the double seismic zone just
under these LF events, again have normal predominant frequen
cies for P-and S-waves. This indicates that the low predominant
frequencies of these anomalous events are not due to the propa
gation path but to the source.
The hypocentres of the LF events are indicated by crosses in
Fig. 1 and by red circles in Fig. 2a. We can see from these
figures that LF events are located approximately beneath active
FIG. 1 Fractional P-wave velocity perturbation (in %) at a depth of 40 km
beneath northeastern Japan. The velocity perturbation is from the mean
value of estimated velocities at this depth. Red and blue colours, low and
high velocities. respectively. Solid triangles. crosses. open squares and open
circles denote the locations of active volcanoes, low-frequency microearth
quakes at depths of 25-40 km. distinct S-wave reflectors in the mid-crust
and large crustal earthquakes. respectively.
NATURE · VOL 352 · 22 AUGUST 1991 ARTICLES
volcanoes or around low- V zones. The magnitudes of these LF
events are not large enough to determine a unique focal mechan
ism solution by using the initial motions of the P-wave. It seems
that the distributions of P-wave initial motions on the focal
sphere for two relatively large events cannot be explained by a
double-couple mechanism, although the polarity data are too
sparse and their coverage on the focal sphere are too poor to
insist on non-double-couple mechanisms for these two events.
Nevertheless the closeness to active volcanoes and other fea
tures described above strongly suggest that these anomalous
events are caused by magmatic activity in the depth range 25-
40km.
Distinct S-wave reflectors
We have detected distinct S-wave reflectors in the mid-crust at
two locations in northeastern Japan: one in northern Akita
prefecture and the other in southeastern Yamagata prefecture.
Figure 4 gives examples of seismograms showing reflected S
waves from one of these two reflectors. We can see a sharp
impulsive later phase following the direct S-wave at all three
stations. The main features of this unusual phase are as follows:
( 1) the phase is most clearly defined on instruments measuring
the horizontal component; (2) the predominant frequency of
the phase is nearly the same as that of direct S-wave; and (3)
the amplitude of this phase is very large, and in some cases
even larger than that of the direct S-wave. Arrival-time analysis
of this phase together with the above-mentioned features indi
cate that this unusual phase is an S-wave reflected from a sharp
velocity discontinuity in the mid-crust (S x S phase), similar to
that first detected and identified beneath the Rio Grande rift
near Socorro, New Mexico23•24• A reflected and S-to-P converted
phase (S x P phase) at the same discontinuity can be also found
at station NIB. This S x P phase is most clearly seen in the
vertical component, although its amplitude is not very large.
The large amplitude of the S x S phase and the ratio of S x P to
S x S amplitudes are explained by a large velocity contrast across
38' N
-6% 0% 6%
P-wave velocity perturbation
685
a ARTICLES
0%
P-wave velocity perturbation
b Volcanoes
~
..c
i5.
Q)
0
20
a disconuuuuy uuu10dain by a very-low-rigidity material, such
as a magma body, as pointed out by Sanford et a/.23.
We try to map this velocity discontinuity by analysing the
arrival times of the S x S phase observed at the four stations
shown in Fig. 5. By assuming the velocity discontinuity to be a
single flat plane, we can locate the reflection point for each of
the observed S x S phases from the S x S arrival-time data in
combination with accurately determined hypocentre coordin
ates. The dip of the discontinuity is corrected in the same way
as seismic migration. Estimated locations of the reflection points
are indicated by open circles in Fig. 5. They form a gently
inclined flat plane with scatter of ± 1 km, showing that the
original asumption that the velocity discontinuity is a simple
flat plane is adequate . The amount of scatter (±1 km) corre
sponds to the accuracy of the hypocentre locations and that of
the S x S phase readings. As presently mapped, this unusual
S-wave reflector is distributed over an area of 15 x 15 km2 at
depths of 12-17 km beneath a Quaternary volcano. The reflector
slopes at an angle of -15?, becoming shallower towards the
ESE, in which direction more active volcanoes are densely
distributed. Beneath this reflector, a group of LF microearth
quakes occur at depths 27-37 km.
Distinct S-wave reflectors, similar to that described above,
686 FIG. 2 a, East-west vertical cross-section of fractional P-wave velocity
perturbation (in %) along the line AB in the inset map. Open circles are
microearthquakes within a 60-km width along AB located by the seismic
network of Tohoku University in 1987-1990. Red circles denote low
frequency microearthquakes located in 1975 -1990, and a red line represents
the location of the mid-crustal S-wave reflector. The land area and active
volcanoes are shown at the top of the figure by the bold horizontal line and
red triangles , respectively . The depth distributions of the Conrad and Moho
discontinuities and the top of the subducted Pacific plate are fixed in the
inversion , and are shown by bold lines . The estimated location of the bottom
of the Pacific plate is also shown by a bold line. b. Vertical cross-section
of fractional P-wave velocity perturbation along the line CD in the inset map.
Velocity perturbation is shown by the colour scale as in a.
have also been found in two other locations in northeastern
Japan25-27• The locations of the reflectors detected here and
those detected in previous work are plotted by large squares in
Fig. 1. As is obvious from this figure, all the reflectors are located
near active volcanoes and/ or in or around the low-V zones. LF
microearthquakes occur at greater depths beneath the reflectors .
These observations and the features of S x S phase described
above support the evidence that the reflectors are the upper
surfaces of magma bodies in the mid-crust. Detection of the
reflector requires a favourable geometrical relation between
source, receiver and reflector. Even if a distinct reflector actually
exists, we cannnot detect it without both earthquakes and
stations just above the reflector. This suggests that many other
reflectors may exist at different locations in northeastern Japan,
because seismicity is not homogeneous in space and seismic
stations are not densely distributed. The velocity decrease of
several percent in the low-V zones estimated by the present
tomographic study also suggests the existence of partially molten
materials in the mid-crust and in the upper mantle.
No reliable information has been obtained to determine the
bottom or the thickness of the magma body hypothesized here.
Direct S-waves at station MOR from intermediate-depth
earthquakes right under it, which pass through the estimated
NATURE · VOL 352 · 22 ALXlUST 1991
FIG. 3 Examples of short-period (1 s) three-component seismograms of a,
an unusually deep (32 km) low-frequency microearthquake and b, a normal
depth (10 km) event, which occurred near lwate volcano. Magnitudes of the
two events are 1.6 and 2.8, respectively.
magma body in the mid-crust, have spectra no different from
those at other stations. In other words, S-waves passing through
the magma body are not significantly attenuated. This fact
suggests that the magma body is very thin, probably between
several hundred metres and a few kilometres, similar to that
detected beneath the Rio Grande rift in the vicinity of Socorro,
New Mexico24•
Deep structure
In the land area of northeastern Japan, shallow crustal seismicity
is confined to the upper 15 km of the crust, with the exception
of a few deep LF microearthquake s as already shown. If
examined in more detail, the greatest depth of this shallow
seismicity, which is sharply delimited, varies with the location.
We have found that, in the central part of northeastern Japan
where seismic stations are densely distributed, the cut-off depth
for the shallow seismicity locally elevated becomes shallower
at some places near active volcanoes ; this can be explained by
the local elevation of the brittle-ductile transition zone because
of the higher temperature in the crust in those areas.
The crust and upper-mantle structure beneath northeastern
Japan can be inferred as schematically illustrated in Fig. 6, by
combining all the observations presently obtained : low-V zones
in the crust and upper mantle, LF microearthquakes at depths
25-40 km, S-wave reflectors in the mid-crust and local elevation
of the base of the seismogenic zone beneath active volcanoes .
Low-V zones in the uppermost mantle are the manifestation of
mantle diapirs which form the roots of arc volcanoes at depth.
Magmatic activity of the mantle diapir may cause the occurrence
of LF microearthquakes around it, although they occur very
infrequently. In the mid-crust, molten materials rising from
NATURE · VOL 352 · 22 AUGUST 1991 ~ EW :::;:
I
UD
p s
h
' EW
~ ARTICLES
below show themselves as distinct S-wave reflectors which are
presumed to be the tops of very thin magma bodies.
Shimamoto28 estimated the spatial distribution of mechanical
strength in the crust and upper mantle beneath the northeastern
Japan arc from the rheological properties of rocks. Temperature
increases relatively rapidly with depth beneath this volcanic arc;
this results in a very thin brittle seismogenic zone in the crust.
Its thickness is estimated to be -15-20 km, corresponding to
the actual cut-off depth of 15 km for shallow crustal seismicity .
The lower portion of the crust and the wedge mantle, below the
base of the brittle seismogenic zone, are governed by creep or
flow, being weak and incapable of supporting much stress.
Consequently horizontal compressional stress in the direction
of relative plate motion, acting on the crust and the upper mantle,
is supported mostly by the upper 15 km of the crust, the relatively
strong seismogenic zone.
In this situation, the stress will concentrate in or around the
places where the base of the seismogenic zone is locally elevated,
and finally would cause shallow crustal earthquakes. In places
where the base of the seismogenic zone is extremely shallow,
such as in the vicinity of active volcanoes, the stress will be
released by creep or flow, and thus shallow crustal earthquakes
will occur not in but around such areas. Local elevation of the
base of the seismogenic zone may be due to molten materials
having ascended from greater depths. Shallow crustal earth
quakes are not homogeneously distributed in space over the
land area beneath this volcanic arc, but are concentrated in
particular locations. We suggest the mechanism described above
is the main cause for this concentration of crustal events.
The depth to the brittle-ductile transition zone is prescribed
principally by the thermal structure in the mid-crust. At present,
687
ARTICLES
KMH <v> NIB (Hl
P, ~ ,sxs P. sxs
I
'-----' 0 5s '------' 0 5s 0 5s
FIG. 4 Examples of short-period seismograms of shallow microearthquakes
that occurred near Mt Moriyoshi, showing clear later arrivals of S-waves
which are presumed to be reflected S-waves (S xS) at the top of magma
body. Arrival times of direct P-and S-waves are denoted by P and S on the
top, and V and H in parentheses indicate vertical and horizontal components.
Locations of seismic stations MOR. KMH and NIB are shown by crosses in
Fig. 5, where the focal area of these events is indicated by the shaded area.
the spatial distribution of temperature in the mid-crust is not
precisely known. As a substitute for the temperature distribution,
we use the P-wave velocity distribution at a depth of 40 km;
this is the depth at which the most accurate velocity distribution
is obtained, according to the resolution analyses. If low velocity
corresponds to high temperature, we can consider Fig. 1 as a
map showing a first approximation to the temperature distribu
tion. Thus the base of the seismogenic zone might be locally
elevated in the low-V zones in Fig. 1 because of the higher
temperature, causing stress concentration around these zones
under the tectonic stress field of horizontal compression. In Fig.
1, open circles indicate large crustal earthquakes that occurred
since 1931 (ref. 29). It is seen from this figure that many of these
occurred around the low-V zones. Crustal earthquakes with
smaller magnitudes have a similar tendency, which is partly
seen on the vertical cross-section in Fig. 2a; many events in the
upper crust beneath the land area are concentrated in or around
the low-V zones. The low-V zone corresponding to high seis
micity is also found in the Mount St Helens region and is
presumed to be the lineation of a crustal weak zone30• Of course,
the P-wave velocity distribution obtained in this study is not an
adequate approximation for the temperature distribution in the
crust, which is the information we truly need to know precisely.
Although Fig. 1 seems partly to support our interpretation, ~ NSK
B A NIB MOR .&
o~----------~~~----~~--~
earthquake swarm
_oo~ ~o •• O
-o reflector
.r. 20 a
Q)
Q
30 0
0
0 0
o 'b\lo
8 .1'. low-frequency
s'tlol'i mlcroearthquake
4QL-------------------------~
FIG. 5 Locations of deep low-frequency microearthquakes, the mid-crustal
S-wave reflector and an earthquake swarm beneath Mt Moriyoshi, northern
Akita Prefecture. a, Epicentres of low-frequency events (open squares),
earthquake swarm area (shaded area) and S-wave reflection points (open
circles). b, Vertical cross-section along AB in a. Solid triangle and crosses
show the locations of Mt Moriyoshi (a Quaternary volcano) and observation
stations, respectively.
/ .......
transition '- -reflector ..- reflector
[) lower crust Low-V Low-V (J FIG. 6 Schematic east -west cross-section of the
crust and upper-mantle structure beneath north
eastern Japan. The depth to the brittle-ductile
transition zone in the crust locally becomes shal
low, perhaps because of ascending magma, which
will cause stress concentrations around regions
where the base of the brittle seismogenic zone is
locally elevated. ----~~-ilJ_LL~~~~~cy--~1~} __ _ * microearthquake * Low-V
upper mantle Low-V
688 NATURE · VOL 352 · 22 AUGUST 1991
further investigation of the detailed structure of the crust and
upper mantle is necessary for confirmation.
We can say, however, that the observations obtained here
from seismic data indicate the deep structure of volcanoes
beneath the northeastern Japan arc. Deep roots of volcanoes
Received 19 March; accepted 23 July 1991.
1. Yoshii, T. Tectonophys. 55, 349-360 (1979).
2. Hasegawa, A .. Umino, N. & Takagi, A. Tectonophys. 47, 43-58 (1978).
3. Engdahl, E. R. & Scholz, C. H. Geophys. Res. Lett. 4, 473-476 (1977).
4. Takagi, A .. Hasegawa, A. & Umino, N. J Phys. Earth 25, S95-104 (1977).
5. Hirahara, K .. lkami, A .. Ishida, M. & Mikumo, T. Tectonophys. 163, 63-73 (1989).
6. Hasemi, A. H., Ishii, H. & Takagi, A. Tectonophys. 101, 245-265 (1984).
7. Obara, K., Hasegawa, A. & Takagi, A. Zisin 39, 201-215 (1986).
8. Yoshii, T. & Asano, S. J Phys. Earth, 20, 4 7-57 (1972).
9. Horiuchi, S. et at. J Phys. Earth 30, 71-86 (1982).
10. Zhao, D., Horiuchi, S. & Hasegawa, A. Tectonophys. 181, 135-149 (1990).
11. Matsuzawa, T., Kono, T., Hasegawa, A. & Takagi, A. Tectonophys. 181, 123-133 (1990).
12. Zhao, D .. Horiuchi, S. & Hasegawa, A. Abstr. Ann. Meeting Seism. Soc. Japan No. 2, 86 (1989).
13. Um, J. & Thurber. C. H. Bull. Seism. Soc. Am. 77, 972-986 (1987).
14. Paige, C. C. & Saunders, M. A. ACM Trans. Math. Software 8, 195-209 (1982).
15. Yoshii, T., Bull. Earthquake Res. lnst. Univ. Tokyo 54, 75-117 (1979).
16. Hasegawa, A. et at. Zisin 36, 129-150 (1983).
17. Humphreys, A. & Clayton, R. W. J geophys. Res. 93, 1073-1083 (1988). ARTICLES
are seen from the clear low-V zones of P-wave in the crust and
the wedge mantle beneath active volcanoes, the anomalously
deep (25-40 km) LF microearthquakes around the low-V zones,
and the distinct mid-crustal reflectors of S-wave in or around
the low-V zones. D
18. Inoue, H., Fukao. Y., Tanabe, K. & Ogata, Y. Phys. Earth planet. Int. 59, 294-328 (1990).
19. Umino, N., Matsuzawa, T. & Hasegawa, A. Abstr. Ann. Meeting Seism. Soc. Japan No. 2, 200
(1990).
20. Sibson, R. H. J. geo/. Soc. London 133, 191-214 (1977).
21. Chen, W. & Molner, P. J geophys. Res. 88, 4183-4214 (1983).
22. Tse, S. T. & Rice, J. R. J geophys. Res. 91, 9452-9472 (1986).
23. Sanford, A. R .. Alptekin, 0. & Toppozada, T. R. Bull. Seism. Soc. Am. 63, 2021-2034 (1973).
24. Ake, J. P. & Sanford, A. R. Bull. Seism. Soc. Am. 78, 1335-1359 (1988).
25. Mizoue, M., Nakamura, I. & Yokota, T. Bull. Earthquake Res. lnst. Univ. Tokyo 57, 653-886 (1982).
26. lwase, R. et at. Abstr. Ann. Meeting Seism. Soc. Japan No. 1, 185 (1989).
27. Horiuchi, S. eta/. Tohoku geophys. J. 31, 43-55 (1988).
28. Shimamoto, T. in Comprehensive Rock Engineering (Pergamon. 1991).
29. Utsu, T. Bull. Earthquake Res. /nst. Univ. Tokyo 57, 401-463 (1982).
30. Lees, J. M. & Crosson, R. S. J geophys. Res. 94, 5716-5728 (1989).
ACKNOWLEDGEMENTS. We thank A. T. Linde at DTM for critically reading the manuscript. In the
tomographic study we used arrival time data from the Japan University Network Earthquake Catalogue
published by Earthquake Research Institute, University of Tokyo.
Cloning of a human gene encoding the general
transcription initiation factor liB
llho Ha, William S. Lane* & Danny Reinbergt
Department of Biochemistry, Robert Wood Johnson Medical School, University of Medicine and Dentistry of New Jersey, 675 Hoes Lane, Piscataway,
New Jersey 08854-5635, USA
*Harvard Microchemistry Facility, Harvard University, Cambridge, Massachusetts, USA
Transcription factor liB (TFIIB) has a central role in
transcription of class II genes. The purification of
the human TFIIB protein and isolation of a com
plementary DNA encoding TFIIB activity is reported
here. The sequence of TFIIB, which seems to be
encoded by a single gene, contains a repeated
motif, in addition to a motif with similarity to the
prokaryotic u-factors. The recombinant protein
expressed in bacteria substituted for all the func
tions attributed to the human TFIIB protein.
INITIATION of transcription at class II promoters requires, in
addition to RNA polymerase II, two functionally distinct classes
of transcription factors. One such class, the general transcription
factors, operates through common promoter elements (TATA
and/ or initiator motifs) and is required for transcription of all
class II genes analysed. At least seven different protein factors:
TFIIA, -liB, -liD, -liE, -IIF, -IIG and -IIH comprise the general
transcription factors1•2•3• Binding of TFIID to the TATA motif
seems to be the first step in the formation of a transcription
competent complex, providing a recognition site for the associ
ation of the other general transcription factors and RNA poly
merase II (refs 4, 5, 50, 51). The gene encoding human TFIID
has been isolated6-8. The activity is in a single polypeptide of
relative molecular mass about 37,000 (M" 37K). But TFIID
isolated from He La cells purifies as a large complex (larger than
150K), presumably because TFIID is associated with other
undefined proteins9• Kinetic analyses and complex formation
t To whom correspondence should be addressed.
NATURE · VOL 352 · 22 AUGUST 1991 studies have suggested that TFIIA acts at an early step of the
transcription cycle10-12, presumably by facilitating binding of
TFIID to the TATA box, thereby generating the TFIID-IIA
(DA) complex4•5• TFIIA has been purified to apparent
homogeneity using affinity chromatography on columns contain
ing immobilized TFIID protein (P. Cortes & D.R., manuscript
in preparation). TFIIB associates with the DA complex resulting
in the formation of the TFIID-IIA-IIB (DAB) complex5• The
DAB complex is then recognized by RNA polymerase II. The
association of the polymerase requires TFIIF, a factor that
interacts with RNA polymerase II (refs 13, 14), and is composed
of two subunits15, the smallest of which contains sequence
similarity to the prokaryotic factor, sigma 70 (ref. 16). The
association of TFIIF/RNA polymerase II with the DAB com
plex allows recognition by TFIIE, TFIIG and TFIIH which
bind to the complex (0. Flores et al., manuscript in preparation).
TFIIE has been purified to apparent homogeneity and co purifies
with polypeptides of34 and 56K (refs 17, 18). TFIIG and TFIIH
are two newly identified general transcription factors that have
been partially purified (ref. 3, 0. Flores et al., manuscript in
preparation).
The second class of transcription factors recognizes specific
DNA sequences present in promoters19• These factors, the
specific transcription factors, regulate levels of expression and
in some cases mediate the response to a specific stimulus. The
molecular mechanism(s) by which the specific transcription
factors affect transcription is unknown. But there is evidence to
suggest that these factors communicate directly (through pro
tein-protein interactions) or indirectly (DNA mediated) with
the general transcription factors. Recent studies have suggested
that the activating domain of VP16 can directly interact with
TFIID (ref. 20). It has also been suggested that the VP16
activator can facilitate the association ofTFIIB with the preiniti
ation complex21• Towards our goal of unequivocally determining
689
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Hasegawa (1991) deep structure of northeastern japan.txt
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JOURNAL OF THE GEOLOGICAL SOCIETY OF JAPAN, Vol. 89, No.12, p. 731-732, December, 1963 OPHIOLITE INTHE KAMUIKOTANZONE HOKKAIDO,JAPAN HIDEO ISHIZUKA*,MAKOTo OKAMURA*and YAsUJ1 SAITO** The Horokanai ophiolite,occurring within the Kamuikotan zone,about 30 km northwest ofAsahikawa, central Hokkaido, is superposed in the following general sequence, starting from the bottom and working up:ultramafic rocks with minor mafic cumulates,mafic rocks, andchert（AsAHINA& KOMATSU,1979; IsHIZUKA,1980). Recently,from its chert member,we have foundradiolarians available on age determination,which will be briefly described in this short note. The chert member,which is less than 0.1 per cent in relation to the apparent exposure 142°08' of the-ophiolite,can be divided into threc 44°05 types :(1) massive chert, (2)bedded chert, and (3)basaltic fragment-bearing chert.The type(1)and (2)cherts overlie thebasaltic pillow sequence at the top of the ophiolite, while the type (3)chert is intercalated or closely associated with pillow basalts.Radio- larians are abundant in the type (1) and (2) cherts, but less common in the type (3) one. The radiolarian chert treated here is rcd- River- dish in color and belongs to the typc (1) Uryu chert,of which the locality is shown in Fig. 1.1 Under the scanning electron microscope (SEM) and light microscope (LM),a total ofeight genera and fourspeciesofradio- ： larians were identified.Plate I shows SEM and LM photographs of representative radio- Loc larians. Among thc species,Paruicingula hsui has Fig.1.Geologic sketch map of the Horokanai ophiolite (modified after IsHIzukA,1980). characteristic pore pattern of postabdominal chambers, which is common in Zone 2B to 1 :Kamuikotan metamorphic rocks, 2 4 :Horokanai ophiolite (2: chert,3: Zone 3(carly Tithonian）from the chert mafic rocks, 4:ultramafic rocks),5: Received May 16, 1983 Cretaceous Yezo Group,6:Neogene vol- * Department of Geology,Kochi University, Kochi, canic rocks,7:Locality of the radiolarian 780 Japan. chert treated here (Lat.44°04'12'’and **National Science Museum, Tokyo, 160 Japan Long. 142°13′19""). 731 NII-Electronic Library Service 732 Short Notes 1983--12 of the Coast Range Ophiolite, California References (PEssAGNO, 1977). According to BAUMGARTER et al. (1980), Mirifusus mediodilatatus is a junior AIDA, Y., 1982 : Jurassic radiolarian biostratigra- synonym of M. baileyi PEssAGNo (1977), which phy in the Irazuyama district, Kochi Prefecture, is widely recognized in Upper Jurassic strata. Japan. -A preliminary report-. Proceedingsof Praeconocaryomma magnimamma is also known to the First Japanese Radiolarian Symposium, 1981, range from late Jurassic, but is assumed to 1-16.* have its maximum development during Zone ASAHINA, T. and KoMATSU, M., 1979 : The Horokanai ophiolitic complex in the Kamuikotan 2A (early Tithonian) by PESSAGNO (1977). Tricolocapsa sp. and Zhamoidellum are tectonic belt, Hokkaido, Japan. Jour. Geol. Soc. sp. commonly found in the late Jurassic chert Japan, 85, 317-330. from the Sanbosan Group in Shikoku (AIDA, BAUMGARTER, P. O., De WEvER, P. and KoCHER, 1982). D., 1980 : Correlation of Tethyan Late Jurassic- From the present study, it is suggested that Early Cretaceous radiolarian events. Cachiers the chert member of the Horokanai ophiolite de micropaleontologie, 2, 23-72. is correlated to late Jurassic, most probably IsHIzUKA, H., 1980 : Geology of the Horokanai early Tithonian in age. ophiolite in the Kamuikotan tectonic belt, Finally, we wish 1 to express our sincere Hokkaido, Japan. Jour. Geol. Soc. Japan, 86, thanks to Mr. T. HIRAJuMA for his help 119-134.* in the field work, and to Dr. A. TAIRA PEssAGNo, E. A. Jr., 1977 : Upper Jurassic radio- for his critical reading of this manuscript. larian biostratigraphy of the California Coast We would especially like to thank Dr. K. Ranges. Micropaleontology, 23, 56-113. NAKAsEKo for his helpful comments. *in Japanese with English abstract. Explanation of Plate I Scanning electron micrographs (Figs. 6--10) and light micrographs (Figs. 1—5) of representa- tive radiolarians obtained from the chert (Sample No. HO--810625A). Scale bar=200 micron (Figs. 1—-9), 100 micron (Fig. 10). Figs. 1 and 6 : Hsum sp. indet. Fig. 2 : Mirifusus mediodilatatus (Rusr). Fig. 3 : Eucyridium (?)ptyctum ? RiEDEL and SANFILIppo. Fig. 4 : Zhamoidellum sp. indet. Fig. 5 : Praeconocaryomma magnimamma (Risr). Fig. 7 : Tricolocapsa sp. indet, Fig. 8 : Archaeospongoprum sp. indet. Figs. 9 and 10 : Parvicingula hsui PEssAGNO. NII-Electronic Library Service H. IsHIZUKA, M. OKAMURA and Y. SAITO : Plate 1 Jour. Geol. Soc. Japan. Vol. 89, No. 12. December, 1983 NII-Electronic Library Service
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Ishizuka (1983) - Latest Jurassic radiolarians from the Horokanai ophiolite.txt
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The Zsland Arc (1997) 6, 386495
Research Article
Origin of a magnetic lineation on Kyushu Island, Japan
YASUKUNI OKUBO,~ HIROSHI KANAYA,~ AKITSURA SHIBUYA~ AND KIMIO OKUMURA~
lGeologica1 Survey of Japan, 1-1-3 Higashi, Tsukuba, Japan and 2Sumiko Consultants Co. Ltd.,
2-16-9 Kabuki-cho, Shinjuku-ku, Tokyo, Japan
Abstract Several linear magnetic anomalies over continental crust have been identified in
and around the Japanese Islands. The anomalies are probably related to island arc tectonic
structures, but identifying specific sources has been difficult. Several deep holes were drilled
in and around Aso caldera, where a linear anomaly occurs along an active fault. One drillhole
located on the linear anomaly encountered a zone of highly magnetized and altered basement
rocks at least 100 m thick at a depth of -1000 m. The other hole was located away from the
anomaly and did not encounter any high-magnetic zones. Rocks from the zone have excep-
tionally strong remanent magnetization (several tens of A/m) sub-parallel to the present
field. AF demagnetization experiments indicated that the magnetization is hard and stable.
Magnetic modeling indicates that the linear anomaly is caused mainly by this layer. Micro-
scopic examination of core samples shows that the highly magnetized zone includes sec-
ondary magnetic minerals and abundant hydrothermal alterations. Temperatures deter-
mined by fluid inclusions and down-hole temperatures show that the temperature of the
highly magnetized zone was elevated in the past relative to surrounding rocks. The high tem-
perature could destroy primary magnetic minerals and replace them with secondary mag-
netic minerals. Thus, the past hydrothermal system may have enhanced thermo-chemical
remanent magnetization. The results can produce a model indicating that there was a past
hydrothermal system related to the tectonic structure.
Key words: active fault, alteration, Aso caldera, magnetic survey, magnetization, suscepti-
bility, tectonic structure.
INTRODUCTION
Several linear magnetic anomalies have been iden-
tified over both continental and oceanic crust in
and around the Japanese Islands (Fig. l(a))
(Makino et al. 1992a; Okubo et al. 1994). Oceanic
lineations are attributed to sea-floor spreading in
accordance with the Vine-Matthews hypothesis
(Vine & Matthews 1963), but continental lineations
are still a matter of discussion and considerable
interest. They could provide insights on tectonic
evolution because magnetic sources are often
related to tectonic events. For example, thermal
convection along fault zones may produce concen-
trations of magnetic material. Tectonic stresses
could generate uplift of magnetic bodies from
Accepted for publication July 1997. depth. Serpentinization along fault zones is a well-
known example of magnetic mineral concentration.
Moreover, tectonic events produce various
forms of remanent magnetization. Thermal rema-
nent magnetization acquired by temperature
changes and chemical magnetization acquired by
physiochemical changes, as examples, can occur in
tectonically active regions, such as continental
margins. Thus, investigations to identify the
source of magnetic anomalies and the origin of
magnetic material could lead to an improved
understanding of tectonic evolution.
The manner in which magnetic material is pro-
duced is crucial to understanding tectonic struc-
ture and tectonic evolution. These are difficult
questions to resolve in Japan. Some lineations in
Japan lie parallel to trenches, whereas others lie
Magnetic lineation on Kyushu Island 387
gests that the anomaly is related to the tectonic
evolution of the island arc and may reflect static
and dynamic crustal structures (Okubo et al. 1994).
But evidence of the origin has not yet been
revealed.
In this paper, we discuss the origin of magnetic
lineation “B” (Fig. l(a)) located on Kyushu Island.
Okubo and Shibuya (1993) used gravity and mag-
netic analyses to determine the structure beneath
Aso caldera and concluded that highly magnetized
rocks (>lo A/m) must reside in the Futagawa
fault. Here, we add crucial new evidence. Namely,
highly magnetized rocks were found in a deep
drillhole located on this positive magnetic anomaly.
Analyses of these rocks suggest the origin of the
anomaly and provide clues about the cause of the
lineations in the continental crust. Identification
of the sources allows us to speculate on the origin
of the magnetic lineations in the context of
island arc evolution and illustrates the usefulness
of magnetic lineations for construction of tectonic
models.
10 20 30 50 km
I
Fig. 1 Magnetic lineations of the Japanese Islands and adjoining areas (a) and
magnetic anomaly map of the Aso caldera and vicinity (b). (a) Lineations are
drawn from the magnetic map of the Japanese Islands and adjoining areas
(Makino eta/. 1992a). The small square indicates the location of Figure l(b). (b)
Line A-B shows the location of the profile of Figs 2 and 6. Contour interval is 50
nT, with gray area indicating positive anomaly values. The maps express total
intensity anomalies at 1981 m above sea level. Data are from the Geological
Survey of Japan and the New Energy and Industrial Technology Development
Organization.
along tectonic structures, such as trenches and tee.-
tonic lines (Makino et al. 1992a; Geological Survey
of Japan and CCOP et al. 1993). Ogawa and Suya-
ma (1975), Segawa and Oshima (1975), Honkura et
al. (1991), Finn (1994), and Okubo and Matsunaga
(1994) for example, have discussed the tectonic
framework based on the possible sources of mag-
netic lineation “A” (Fig. l(a)) located along the
east side of the Tohoku arc. The alignment sug- LINEAR MAGNETIC ANOMALY ALONG THE FUTAGAWA
FAULT
A positive linear magnetic anomaly lies south of
the Futagawa fault, an active fault extending
west-southwestward from Aso volcano located in
central Kyushu, western Japan (Fig. l(b)) (Makino
et al. 1992b). The anomaly extends parallel to the
fault with an amplitude exceeding 500 nT.
The dominant exposures in this area are vol-
canic rocks which erupted before and during the
formation of Aso volcano. The activity of Aso vol-
cano commenced -300 000 years ago. Pre-Aso vol-
canic rocks, the formation before Aso volcanic
rocks, are production during the activity of Aso
volcano. The current caldera shape was formed by
four large scale eruptions of pyroclastic flows
about 70 000-80 000 years ago and produced Aso
volcanic rocks (Ono & Watanabe 1985).
These volcanic rocks are strongly magnetic;
sample measurements indicate an average of
5A/m (Okubo & Shibuya 1993), but cause only a
part of the observed anomaly. The magnetic effect
of the volcanic rocks can be estimated from model
studies. For example, a uniformly magnetized
topographic model with a magnetization of 5 AJm
has been shown to cause some anomaly (topo-
graphic effect) (Fig. 2). The previous result of
two-dimensional magnetic modeling using a
north-south profile of the observed minus the
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388 Y. Okubo et al.
CORE SAMPLES FROM DRILL HOLES (Obsemd)
- (Topographic Effect) 2 v 750t
Dis .- t-
-2501 ...... ' ....
% v 1000
500 5 .- Y
Fig. 2 Obseved anomalies and observed anomaly minus topographic effects
(Top), and topography (Bottom). See Fig. l(b) for location of profile. A total
magnetization of 5 A/m and a magnetization parallel to the present field (decli-
nation = 8" northwest, inclination = 46" down) are assumed for the topographic
model.
topographic effect about 5 km west of A-B profile
(Fig. l(b)) delineates a highly magnetic prism-
shaped buried body (>lo Mm), dipping northward
at an inclination of 55" (Okubo & Shibuya 1993).
Consequently, buried sources must also contribute
to the magnetic anomaly. Several deep holes have been drilled recently in
and around Aso caldera. Two drillholes reached
basement level but encountered very different
magnetizations. Drillhole AS4 contacted base-
ment at 880 m below the surface. This drillhole is
sited on the linear magnetic anomaly, and altered
and highly magnetized basement rocks were
found. The average value of susceptibility (22 sam-
ples) was 2.6 X lo3 X 4 .n (SI). Microscopic exami-
nations proved that rocks sampled in AS4 are
mafic and bear a relatively high portion of colored
minerals.
Drillhole AS-5, on the other hand, encountered
basement at 950 m below the surface. This drill-
hole is away from the anomaly, and the average
value of basement susceptibility (10 samples) was
only 0.016 X lo3 X 4 .n (SI) (Table 1).
To study susceptibility, remanence and modal
amounts of colored minerals of the basement
rocks of AS4, 15 core samples were selected and
examined.
Table 1 Susceptibility from core samples of AS4 and AS-5.
AS4 AS-5
Formation Depth (m) Susceptibility K X lo3 X 4 n(S1) Depth (m) Susceptibility K X lo-" X 471 (SI)
Pre-Aso volcanics
400.5
600.5
800.0
Dike801.0 2.565 840.0
860.0
872.5
880.0
900.0
920.0
930.0
940.0
960.0
980.5
1000.0
1020.0
1040.0
1060.0
1080.0
1090.0
1100.0
1102.0
1120.0
1140.0
1160.0
1180.0
1182.5
1201.0 Basement rocks 1.783
3.543
1.093
4.146
4.162
4.658
0.013
1.861
3.456
0.509
1.241
0.154
0.808
5.245
3.963
1.007
2.605
3.064
2.741
2.849
3.319
3.182
6.580
0.015
0.064
3.034
6.852 400.0
600.0
768.0
801.0
980.0
1000.0
1040.0 1080.0
1096.7
1100.0
1120.0
1140.0
1160.0
1200.0 2.512
1.240
0.044
2.565
0.021
0.022
0.012 0.020
0.015
0.013
0.017
0.009
0.019
0.012
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Magnetic lineation on Kyushu Island 389
Inclination
Magnetization (Ah) (deg.)
0 10 20 30 40 50 60 40 50 60 70 Formation 7-
900
1000
1100
m
1200 h iz 1000-
0" r, h
1100-
1200- Modal Amounts of
Colored Mineral
(%)
01234567 7
900// Modal Amounts
Susceptibility
IIIIIII
01234567
Susceptibility
(lx10%4x (SI))
Fig. 3 Comparison of susceptibility and modal amounts of colored mineral of
cores sampled from drillhole AS-4.
Modal amounts of colored minerals were deter-
mined by point-counting using polished faces of
each core sample in AS4 (Fig. 3). The mode corre-
lates reasonably well with the susceptibility, and
suggests that susceptibility depends on the
amount of colored minerals. Susceptibility reflects
variations in the abundance of ferrimagnetic min-
erals and 1% magnetite by volume corresponds to
a susceptibility of 2.6-3.0 X 47t (SI) (Clark &
Emerson 1991). This relationship holds for the
interval of 1050-1100 m, but susceptibilities else-
where are generally too low for magnetite. This
implies that much of the colored mineral must be
magnetite between 1050-1100 m and elsewhere
must be other ferrimagnetic minerals such as
hematite, titanohematite and sulfides.
The rocks of AS-5 are fresh, very felsic, and
contain few colored minerals. We conclude that the
large difference in basement susceptibility stems
from the quantity of ferrimagnetic minerals.
Although basement susceptibility can be quite
high, the linear positive anomaly does not originate
entirely from magnetization induced by the earth's
ambient magnetic field. Since the earth's magnetic
field is less than 50,000 nT, the average intensity of
induced magnetization of the rocks sampled from -
0 10 20 30
Q-Value I Im
Fig. 4 Comparison of induced magnetization, remanent magnetization,
Q-value, inclination of remanent magnetization, and lithology of AS-4. Q-value
is the ratio of remanent magnetization to induced magnetization. Induced mag-
netizations are calculated from susceptibilities multipled by the ambient earth's
magnetic field (47000 nT). Dashed line indicates the present field inclination of
46" down.
AS4 is less than 2 Alm, the same order of magni-
tude as the Aso volcanic rocks covering the area
and is much smaller than the magnetization
inferred from the previous magnetic analysis (>lo
Alm) (Okubo & Shibuya 1993).
Remanent magnetization of cores sampled from
AS4 ranged from less than 1 Alm to 50 Alrn
(Fig. 4). Samples with remanent magnetization
higher than 2 Alm were located at depths between
900-1080 m and had Q-values (remanent magneti-
zatiodinduced magnetization) greater than 1.6.
The interval of 1000-1060 m showed extremely
large Q-values, higher than 10. The interval con-
taining remanent magnetization greater than
5 Alm is at least 100 m thick.
Stepwise alternating field (AF) demagnetization
was carried out to examine the effect of secondary
magnetization such as a drilling-induced overprint
and isothermal remanent magnetization acquired
during the exposure to an ambient magnetic field.
The experiment was done on the four samples of
1000, 1020, 1060 and 1080 m, which have high
Q-values, and the one sample of 1140 m, which has
a low Q-value. All samples showed univectorial
behavior as presented in Fig. 5, and no remarkable
difference was recognized between the samples.
Consequently, we concluded that the cores do not
carry such spurious remanence as acquired by
drilling. This implies that the remanent magnetiza-
tions shown in Fig. 4 are hard and stable in situ
components.
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390 Y. Okubo et al.
S: :N
AS-4
1060 m
\
\
0 I
\
1
1
0 I
1 I
I
\ I
I
1 dev.=0.4 A/m
I 1
1 8 E Dn AS-4
1080 m
Fig. 5 Orthogonal plots of progressive AF demagnetization data of the cores at
depths of 1060 rn and 1080 m sampled from AS4 Solid circles denote prolee-
tions of vector endpoints on the horizontal planes; open circles denote projec-
tions of vector endpoints on the vertical planes. The cores were not oriented to
the azimuth. Steps of demagnetization are 0, 5, 10, 20, 25, 30, 40, 50, 60, 80,
100 rnT
The drill cores are azimuthally unoriented, so
inclination of the remanent was measured. Since
the Aso caldera has been active for only 300 000
years, the remanence would have been acquired
during the Brunhes Normal Chron and inclinations would point to the present field inclination. The
inclinations shown in Fig. 4, which are angles with
respect to the horizontal plane, vary between
40-70". The present field inclination in this area is
46", so the inclinations of samples are generally
concordant with the present field. The differences
could reflect tilts of rocks after the remanence
acquisition.
MAGNETIC MODELING
Forward modeling is a method of interpreting the
cause of magnetic anomalies through trial-and-
error attempts to fit observed anomalies. It is well
known that an infinite variety of models can fit the
same anomalies, but independent information can
help constrain this ambiguity. Drillhole data and
geologic mapping provide several initial para-
meters. Drillhole AS4 detected a highly magne-
tized zone about 100 m thick. The Futagawa fault
strikes west-southwest from Aso volcano, dips
north-northwest, and lies along the south margin
of a broad east-west trending graben. These con-
straints demand a magnetic source that is horizon-
tally long and thin, as shown by the model in
Fig. 6.
Remanent magnetization direction is expected
to be sub-parallel to the present magnetic field,
because the zone resides in the pre-Aso volcanic
rocks that were produced during Brunhes Normal
Chron and the inclinations of the rocks range
around the inclination of the present field. At that
time, the direction of magnetization of the source
was assumed to be parallel to the present field
(declination = 8" NW, inclination = 46" down).
(Observed)
- (Topograplc Effect)
I t .n
Topography 2 sea level
B
P Magnetic Source c1 5 34 Alm
Fig. 6 Magnetic model of Futagawa Fault. Profiles at top are calculated anom-
aly and observed anomaly (minus topographic effects). See Fig. l(b) for location
of profile. A total magnetization of 34 A/m and a magnetization parallel to the
present field (declination = 8" northwest, inclination = 46" down) are assumed
for the model.
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Magnetic lineation on Kyushu Island 391
Since the profile line of “A-B” in Figure l(b) is
perpendicular to the length of the positive mag-
netic lineation, the strike of the model in Fig. 6 is
parallel to the Futagawa fault and the magnetic
lineation. The model indicates that the top of
the source is located south of the Futagawa fault.
The depth to its top face varies from about sea
level to 4 km, and it dips 20” north-northwest. The
model predicts that the body lies at 1000 m below
the surface at the location of drillhole AS-4. In a
magnetic inversion problem such as this, it is diffi-
cult to resolve both the magnitude of magnetiza-
tion and the thickness of the body. Increasing the
thickness of the body requires reducing the mag-
netization, so that these parameters cannot be
uniquely determined. If magnetization is 34 AJm,
the average thickness of the body is 275 m. If mag-
netization is 10 Nm, the average thickness of the
body is greater than 800 m. The thickness of the
highly magnetized zone suggests that a thinner
layer is preferable and, therefore, we prefer a
model of 275 m in thickness and a magnetization of
34 AJm.
GRAIN SIZE
We have investigated the relationships of the
remanent magnetization and grain size of colored
mineral through microscopic examinations of sam-
ples from drillhole AS4. Colored minerals less
than 10 pm are invisible, so the minerals for the
examinations were greater than 10 pm in dia-
meter.
Grain size and microstructure of the magnetic
mineral can influence the magnitude of remanent
magnetization. In general, smaller magnetite
grains acquire stronger thermal remanent magne-
tization (Nagata 1961; Day 1977). The magnetic
properties of a ferro- or ferrimagnetic mineral are
fundamentally controlled by the domain structure
of the mineral grains. The natural remanent mag-
netization of relatively coarse-grained mineral is
often multicomponent in character and represents
the vector sum of remanence components. The
rocks which contain predominantly coarse-grained
magnetic minerals yield generally low Q values. On
the other hand, very fine grain possesses a single
domain structure and often exhibits very high Q
values. The critical grain length for the single
domain to multi-domain transition is 0.06 pm for
magnetite and 2.4 pm for titanomagnetite (Clark
1983).
The properties of small multi-domain grains, 0 0 0 2 0 0 m
Depth (rn)
Fig. 7 Comparison of natural log of 1/(L x S), remanent magnetization, and Q-
value of AS-4 as a function of depth. L and S are average long and short diam-
eters of colored mineral grains, respectively, in unit of pm. Shaded pattern indi-
cates the depth range where remanent magnetization intensities and Q-values are
high
commonly called pseudo-single domain grains, are
intermediate between those of single domain and
large multi-domain grains and vary systematically
with grain size. The most important property of
pseudo-single domain is the capacity to retain rela-
tive intense, hard and stable remanence analogous
to that of single domain grains (Clark 1983).
Grains less than 10 pm in dimension are invisible
under microscopic examination. Therefore, we
expect that a microscopic examination could not
discover this relationship. In fact, the product of
the average long dimension (L) times the average
short dimension (S) determined petrographically,
which reflects grain size, does not correlate with
the magnitude of remanent magnetization (Fig. 7).
This suggests that the remanence carriers are too
small to be seen petrographically.
THERMAL CHANGE
The history of heating and cooling also affects
remanent magnetization. Thermal remanent mag-
netization is acquired by a magnetic mineral as it
cools from above the Curie temperature in the
presence of a magnetic field (Carmichael 1989).
Chemical magnetization is acquired by a magnetic
mineral as it undergoes certain physicochemical
changes during crystallization (Carmichael 1989).
Temperature change is one of the causes of miner-
al re-crystallization. Thus the heating and cooling
history of a rock has a great influence on the rema-
nent magnetization and often produces thermo-
chemical remanence.
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392 Y. Okubo et al.
Table 2 Lithology and major secondary minerals of cores sampled from drillhole AS4
Depth (m)
840
869.5
890
900
910
950
968
990
1010
1020
1030
1050
1071
1110
1130
1150
1170
1190
1200 Lithology
Andesite lava
Sericite quartz phyllite
Hornblendite cataclasite
Hornblendite cataclasite
Hornblende clinopyroxenite cataclasite
Clinopyroxene gabbro cataclasite
Clinopyroxene gabbro cataclasite
Biotite-hornblende diorite
Hornblende gabbro
Meta-gabbro
Clinopyroxenite cataclasite Clinopyroxenite cataclasite
Clinopyroxene gabbro cataclasite
Dunite
Calcareous ultramafic rock ~~
Serpentinite (?)
Serpentinite
Quartz-bearing clinopyroxene
andesite Peridotite Major secondary minerals
Calcite, Chrolite
Sericite, Calcite, Hematite
Calcite, Chrolite, Limonite
Actinolite, Chrolite
Actinolite, Chrolite
Actinolite, Chrolite
Actinolite
Chrolite
Actinolite, Chrolite, Magnetite
Actinolite, Chrolite, Magnetite, Albite
Actinolite, Chrolite, Calcite, Magnetite
Actinolite Actinolite, Chrolite, Magnetite, Albite, Calcite
Serpentinization, Chrysotile
Serpentization, Chrysotile, Magnetite,
Carbonate minerals
Serpentization, Chrysotile, Magnetite, Carbonate
minerals
Serpernitzation, Chrysotile, Magnetite, Calcite
Dolomite, Sericite, Chrolite, Calcite, Titanite,
Magnetite
Serpentine ~ ~~~~
Shaded pattern indictes the depth range where remanent magnetization intensities and Q-values are high.
The thermal change must be recorded in the
rocks as evidence of alteration and trapping tem-
peratures of fluid inclusion. To investigate the pos-
sibility of alteration, we microscopically re-exam-
ined the basement cores of AS4 at depths
between 840-1200 m (Table 2). There was a clear
difference on the alteration zones. The samples of
upper formation between 840-1071 m indicate low-
temperature thermal alteration. Quartz is a com-
mon secondary mineral in the interval, but is not
shown in Table 2. Other major minerals related to
alterations are calcite, chlorite, and actinolite. Cat-
aclasite, a metamorphic rock produced by crushing
and granulation mainly associated with fractures,
is prevalent between 890-1071 m. The samples of
the lower formation between 1110-1200 m, on the
other hand, are serpentinite of ultramafic rock ori-
gin.
Magnetite was found at depths of 1010, 1020,
1030 and 1071 m and in the serpentinized zone
between 1030-1090 m, suggesting that new mag-
netic materials were produced by the thermal
alteration and serpentinization.
Actinolite, which is a secondary mineral caused
by thermal alteration and whose crystallization
temperature is over 3OO0C, was found in the cores
at depths between 900-1071 m. In the cores shal-
lower than 900 m, on the other hand, no actinolite
was found, suggesting that the past temperature did not exceed 300°C (Hirano, pers. comm.).
Trapping temperatures of fluid inclusions were
determined. A fluid inclusion is a tiny cavity in a
mineral, 1.0-100.0 ym in diameter, often containing
liquid in high pressure, formed by the entrapment
in crystal irregularities of fluid, commonly that
from which the rock crystallized. With tempera-
ture and pressure decreasing, the entrapped liquid
shifts along isochore and turns into a two-phase
liquid-rich inclusion at the homogenization temper-
ature, and then along a vapor-saturated water
curve. A drillhole core sample, that has been
reduced in temperature and pressure, must pos-
sess vapor saturated fluid inclusions. With an arti-
ficial increase in temperature, the pressure in a
fluid inclusion reverses along the vapor-saturated
water curve and the inclusion finally becomes a
one-phase at the homogenization temperature.
This process is visible petrographically. The shift
along isochore can be estimated under the assump-
tion of pore pressure at which the rock crystallized
(Sasada 1989). Thus, the trapping temperatures of
fluid inclusions found in core samples of AS4 were
determined.
Figure 8 compares down-hole temperatures
with the temperature of hydrothermal alteration
as determined by examination of fluid inclusions.
Down-hole temperature increases approximately
constantly with depth, whereas the temperature
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Magnetic lineation on Kyushu Island 393
agreement between calculated and observed
anomalies in Fig. 6 verifies that the anomaly is
caused primarily by the highly magnetized rocks
sampled in AS4 drillhole.
The microscopic examinations suggest that the
highly magnetized zone underwent hydrothermal
alterations as well as tectonic events such as cata-
elastic metamorphism. Actinolite appearance
between 900-1071 m indicates that the past tem-
perature of the zone was higher than the other
zones. Temperatures determined by fluid inclusion
studies indicate that the interval between
945-1020 m encountered high temperatures in the
past relative to surrounding rocks but has subse-
quently cooled to normal down-hole temperatures.
These confirm the occurrence of thermal change in
the highly magnetized zone. Magnetite appearance
between 1010-1071 m suggests that the hydro-
thermal alterations produced secondary magnetic
minerals.
The past high temperature could produce strong
thermo-chemical remanent magnetizations or new
carriers of remanence in the highly magnetic zone
of drillhole AS4 between 900-1080 m. Very high
Q values are acquired by single or pseudo-single
domain magnetic mineral grains which would
be invisible in conventional petrography. The
hydrothermal alteration is expected to have par-
tially replaced the magnetic mineralogy, which is
probably mainly multi-domain magnetite, with sin-
gle or pseudo-single domain secondary magnetic
minerals. Alternatively, the remanence carriers
could be magnetic minerals subdivided by
hematite/magnetite intergrowths. Thus, the very
fine magnetite grains acquired thermo-chemical
remanence. Destruction of primary magnetic min-
erals and their replacement by secondary magnet-
ic minerals could produce relatively small changes
in susceptibility, but substantial increases in rema-
nence.
This finding suggests that the source of the lin-
eation was produced by tectonic or volcanic activi-
ty in the study area. The past thermal structure
revealed by fluid inclusion methods shows a
change in thermal environment in the zone. The
magnetic model indicates a thin, long, north-dip-
ping body lying parallel to the Futagawa fault.
Since a large fault system could produce a wide
permeable layer and often cause hydrothermal
convections in a geothermal area, the geometrical
consistency between the fault system and the mag-
netic source suggests that the fault is related to
the genesis of the highly magnetized zone (Fig. 9a).
At least one origin is possible: a change in the Fig. 8 Temperatures determined from fluid inclusion experiments (shaded
column) on samples from drillhole AS-4. Solid line indicates down-hole tem-
peratures. Fluid inclusion temperature were mapped at every 5°C step. The num-
ber beside each column expresses number of measurements of each step. Open
circle denotes average of the measurement temperature. Shaded pattern indicates
the depth range where remanent magnetization intensities and Q values are high.
determined by fluid inclusion is significantly higher
at any given depth. The differences at 944.6 m and
1019.5 m depth are exceptionally large.
These results suggest that basement rocks
have been in a high-temperature environment in
the past and have cooled to normal temperature;
that is, lower than 100°C. At depths between
900-1071 m, historical temperatures may have
exceeded 200°C.
DISCUSSION
The highly magnetic zone between 900-1080 m in
drillhole AS4 probably reflects some tectonic
event. In fact, cataclasite mainly produced by tec-
tonic activities is prevalent between 890-1071 m.
The zone has both strong and stable remanent
magnetizations greater than 2 AJm with a maxi-
mum of 50 A/m and a high Q-value greater than
1.6.
The linear magnetic anomaly along the Futa-
gawa fault is largely caused by the highly mag-
netic layer sampled in AS4 drillhole. This is
demonstrated by the magnetic model (Fig. 6),
which was constrained by AS4 drillhole in terms
of depth, thickness, and magnitude and direction
of magnetization of the magnetic zone. The
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394 Y. Okubo et al.
I Fault System - (a)
Present
Time
Highly Magnetized Zone
Fault System - (b)
Past
Time
Ancient Thermal Convection System
Fig. 9 A process of generation of a highly magnetized zone. Aso volcano and
the Futagawa fault system have provided the heat source and the pathway,
respectively, forming a hydrothermal convection (b). The highly magnetized zone
appeared while the thermal structure changed to normal down-hole temperatures
(4.
hydrothermal convection system may have taken
place. The convection system requires heat
sources and fluid pathways. Aso volcano and the
Futagawa fault system both are active and could
have provided both the heat source and the path-
ways, respectively, to generate the highly magne-
tized zone (Fig. 9(b)). The possibility shows that
tectonic structure of the island arc played an
important role in the acquisition of the strong
remanent magnetization.
CONCLUSION
A highly magnetized zone was found in a drillhole
located near the Futagawa fault. The zone is about
100 m thick and situated about 1000 m below the
surface. Remanent magnetization, with magni-
tudes ranging from 5 to 50 A/m, forms most of the
total magnetization. AF demagnetization shows
that all of the remanent magnetizations of the sam-
ples are hard and stable.
Direction of the magnetization could be sub-par-
allel to the present magnetic field, because the rocks have been produced for the last 300000
years and the inclinations of the samples scatter in
a small range around the present field.
A magnetic model suggests that the main
source of the magnetic lineation lies south of the
Futagawa fault. The model is consistent with
the highly magnetized zone found in AS4 drill-
hole.
Secondary minerals produced by hydrothermal
alteration such as actinolite and magnetite were
found in the highly magnetized zone by micro-
scopic examination. Temperatures determined
from fluid inclusion experiments and down-hole
temperatures indicate that a great thermal change
affected the highly magnetized zone.
The past high temperature zone corresponds to
the highly magnetized zone, suggesting that the
hydrothermal alterations produced petrographi-
cally an invisible single domain or pseudo-single
domain structure of magnetic grain which has the
capacity to retain relative intense, hard and stable
thermo-chemical remanence.
The hydrothermal system and its change could
be boosted by the regional tectonic and thermal
structures of the study area. Thus, magnetic lin-
eations could indicate the location of tectonic struc-
tures and their interpretation could reveal the his-
tory of tectonic evolution.
ACKNOWLEDGEMENTS
Dr Hideo Hirano, Geological Survey of Japan,
determined the microscopic character of the core
samples. Dr Toshitsugu Yamazaki, Geological Sur-
vey of Japan, examined the alternating field
demagnetization of samples and gave us helpful
suggestions on the results. Dr Masakatsu Sasada,
Geological Survey of Japan, gave us helpful com-
ments on temperature determination from fluid
inclusion. Dr D.A. Clark, CSIRO Division of
Exploration, and Dr Richard J. Blakely, US Geo-
logical Survey, reviewed our paper and gave us
helpful comments. Prof. Yozo Hamano, The Uni-
versity of Tokyo, gave us valuable information
about the possible cause of strong remanent mag-
netization. The Association for the Development of
Earthquake Prediction gave us opportunities for
discussion in the course of several meetings. The
New Energy and Industrial Technology Develop-
ment Organization offered drillhole cores, the data
obtained from the drillholes and the aeromagnetic
data. We would like to express our thanks to all of
them.
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Magnetic lineation on Kyushu Island 395
netic map of the Japanese Islands and its explana-
tory text, 1:2,000,000 Map Series (23). Geological
Survey Japan, 24p. (in Japanese with English
abstract 7p.).
NAGATA T. 1961. Rock magnetism. Maruzen, Tokyo.
OGAWA K. & SUYAMA J. 1975. Distribution of aeromag-
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Availableonline at www.sciencedirect.com TECTONOPHYSICS SCIENC DIRECT ELSEVIER Tectonophysics 388(2004)59-73 www.elsevier.com/locate/tecto Upper and middle crustal deformation of an arc-arc collision across Hokkaido, Japan, inferred from seismic refraction/wide-angle reflection experiments Takaya Iwasakia*, Keiji Adachia, Takeo Moriyab, Hiroki Miyamachic, Takeshi Matsushimad, Kaoru Miyashita, Testsuya Takedaa, Takaaki Tairab, Tomoaki Yamadaa, Kazuo Ohtake? Earthquake Research Institute, University of Tokyo, Yayoi 1-1-1, Bunkyo-ku, Tokyo 113-0031, Japan Graduate School of Science,Hokkaido University, Sapporo 060-0810, Japan Faculty of Science, Kagoshima University, Kagoshima, 890-0065, Japan dInstitute of Seismology and Volcanology, Graduate School of Science,Kyushu University,Shimabara,855-0843,Japan Japan Meteorological Agency, Tokyo 100-8122, Japan Received 11 August 2003; received in revised form 31 March 2004; accepted 13 June 2004 Available online 27 August 2004 Abstract The Hidaka Colision Zone (HCZ), central Hokkaido, Japan, is a good target for studies of crustal evolution and deformation processes associated with an arc-arc collision. The collision of the Kuril Arc (KA) with the Northeast Japan Arc (NJA), which started in the middle Miocene, is considered to be a controlling factor for the formation of the Hidaka Mountains, the westward obduction of middle/lower crustal rocks of the KA (the Hidaka Metamorphic Belt (HMB)) and the development of the foreland fold-and-thrust belt on the NJA side. The “Hokkaido Transect' project undertaken from 1998 to 2000 was a multidisciplinary effort intended to reveal structural heterogeneity across this collision zone by integrated geophysical/geological research including seismic refraction/reflection surveys and earthquake observations. An E-W trending 227 km-long refraction/wide- angle reflection profile found a complicated structural variation from the KA to the NJA across the HCZ. In the east of the HCZ, the hinterland region is covered with 4 4.5 km thick highly undulated Neogene sedimentary layers, beneath which two eastward dipping reflectors were imaged in a depth range of 10-25 km, probably representing the layer boundaries of the obducting middle/lower crust of the KA. The HMB crops out on the westward extension of these reflectors with relatively high Vp (>6.0 km/s) and Vp/Vs (>1.80) consistent with middle/lower crustal rocks. Beneath these reflectors, more flat and westward dipping reflector sequences are situated at the 25-27 km depth, forming a wedge-like geometry. This distribution pattern indicates that the KA crust has been delaminated into more than two segments under our profile. In the western part of the transect, the structure of the fold-and-thrust belt is characterized by a very thick (5-8 km) sedimentary package with a velocity of 2.5-4.8 km/s. This package exhibits one or two velocity reversals in Paleogene sedimentary layers, probably formed by imbrication associated with the collision process. From the horizontal distribution of these velocity reversals and other * Corresponding author. Tel./fax: +81 3 3816 1159. E-mail address: iwasaki@eri.u-tokyo.ac.jp (T. Iwasaki). 0040-1951/S - see front matter @ 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2004.03.025 60 T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 geophysical/geological data, the rate of crustal shortening in this area is estimated to be greater than 3-4 mm/year, which corresponds to 40-50% of the total convergence rate between the NJA and the Eurasian Plate. This means that the fold-and- thrust belt west of the HCZ is absorbing a large amount of crustal deformation associated with plate interaction across Hokkaido Island. 2004 Elsevier B.V. All rights reserved. Keywords: Crustal structure; Arc-arc collision; Wide-angle; Hokkaido 1. Introduction tion of the Pacific Plate beneath the Kuril Trench, the latter being responsible for the southwestward migra- Hokkaido Island, Northern Japan, is located along tion of the Kuril Arc sliver (KA) and its collision into the NW Pacific margin, beneath which the Pacific the NJA (Kimura et al., 1983; Kimura, 1986; Kimura Plate is being subducted westward from the Kuril and Tamaki, 1986, see Fig. 2). High-T metamorphic Trench (Fig. 1). West of this island, there exists a rocks (the Hidaka Metamorphic Belt (HMB)) exposed relatively new plate boundary between the North in the western part of the HB show petrological American (Okhotsk) and Eurasian (Amurian) Plates properties almost equivalent to continental or island oriented almost in a N-S direction (e.g. Nakamura, arc middle/lower crustal rocks (Komatsu et al., 1986). 1983). The crustal evolution of Hokkaido Island has Probably, these rocks represent a deeper crustal part of been dominated by a series of accretion and collision the KA including the HB, which was thrust up along processes occurring from the late Jurassic to the the Hidaka Main Thrust (HMT) due to the collision present (e.g. Maeda, 1986; Niida and Kito, 1986; processes mentioned above (Kimura, 1986, 1990; Sakakibara et al., 1986; Kimura, 1994, Fig. 2). The Kimura and Miyashita, 1986). western part of Hokkaido is the northern extension of The structural variation across Central Hokkaido the NE Japan Arc (NJA), which consists of a late was investigated by refraction/wide-angle reflection Jurassic accretionary complex with Cretaceous gran- experiments in 1984 and 1992 (Research Group for itic intrusion. The central part of Hokkaido is Explosion Seismology (RGES) 1988, 1993; Ozel et composed of almost N-S trending tectonic units al., 1996; Iwasaki et al., 1998; Moriya et al., 1998). (Fig. 2). The Sorachi-Yezo Belt (SYB) is character- However, the data from these works do not have ized by ophiolite sequences, Cretaceous forearc basin enough resolving power to reveal the complicated sedimentary rocks and high-P metamorphosed rocks crustal deformation associated with the above colli- (Kamuikotan Metamorphic Belt (KMB). The Hidaka sion processes. A clear image of the collision structure Belt (HB) was created by the westward subduction of was firstly obtained through a series of reflection the oceanic plate during the early Cretaceous to early profiles undertaken in the southernmost part of the Paleogene. The Yubetsu and Tokoro Belts (YB and HCZ from 1994 to 1997 (Arita et al.,1998; Tsumura TB) are also accretionary complexes developed in the et al., 1999; Ito, 2000, 2002, Fig. 2). These results late Cretaceous arc-trench system, in which fragments show the delamination of the KA crust characterized of Jurassic seamounts were incorporated (Sakakibara by an obducting upper crust 23 km thick along et al., 1986). Eastern Hokkaido represents the south- HMT and a descending lower crust. The multi- western margin of the Kuril Arc including Cretaceous disciplinary project, the “Hokkaido Transect", in volcanic and sedimentary rocks. 1998-2000 was intended to reveal structural hetero- The western part of Central Hokkaido is a highly geneity at various scales across this collision zone by deformed region under a compressional stress regime seismic refraction/reflection surveys and very dense lasting from the Paleogene or Early Miocene time. earthquake observations (Iwasaki et al., 2000a,b, This tectonic framework was controlled by the dextral 2001b,2002,2003a,b; Adachi, 2002; Katsumata et oblique collision between the Eurasian and North al., 2002, 2003). This paper presents an upper and American (Okhotsk) Plates and the oblique subduc- middle crustal structure model obtained from the T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 61 130° 135° 140° 145° N.American Eurasian (Amurian) Plate (Okhotsk) 45° An Plate Hokkaido Hid Collisic ne 40° theSeaofJapan HonshuIs PacificPlate 35 Pacific Philippine Sea Plate Ocean 30° Fig. 1. Tectonic map in and around Japan showing the lithospheric boundaries. The refraction/wide-angle reflection lines of a multidisciplinary project, the Hokkaido Transect are also indicated together with other seismic lines discussed in the present paper. refraction/wide-angle reflection experiments con- and western zones, by the Hidaka Main Thrust (HMT). ducted in 1999-2000. The metamorphic grade in the main zone of the HMB decreases away from the HMT, i.e., granulite facies, amphibolite facies, greenschist facies and non-meta- 2. Tectonic setting morphic sediments represented by theNakanogawa Group (Komatsu et al.,1986). Such features of the The Hidaka Collision Zone (HCZ) is composed of metamorphic and igneous rock sequences are compat- N-S trending accretionary and melange complexes of ible with those of continental or island arc crust. The the Idon'nappu Belt (IB) and the Hidaka Belt (HB) main zone is interpreted to be a deeply eroded crustal (e.g. Kimura, 1994; Kiyokawa, 1992; Fig. 2). In the Hidaka Mountains along the western margin of the petrological properties of the middle/lower crust of the HB, high-T type Paleogene to Miocene metamorphic KA side. The western zone, on the other hand, is rocks are exposed in the Hidaka Metamorphic Belt formed from oceanic crustal rocks (Komatsu et al., (HMB). This belt is separated into two parts, the main 1983, 1986; Komatsu, 1986; Osanai et al., 1986). 62 T. Iwasaki et al./Tectonophysics 388 (2004) 59-73 1410 1420 1430 144°E 45N 450 ThrustFault Quaternary OkhotskSea 1 Japan Sea KMB HB 440 YB 449 TB IB SYB HMB Kuril Forearc KMB NE Japan (KA) 430 Arc(NJA) 2 430 M-5 L-6 L-1A S-2 ? 1999Profile ★ L-1 2000 Profile Pacific Ocean 1994-1997ReflectionLines 420 Hinterland HCZ 420 and-thrustbelt 1410 1420 1430 1440 Fig. 2. Geological map of Central Hokkaido with our seismic refraction/wide-angle reflection profiles and shot points (stars). Seismic reflection lines of the Hokkaido Transect were laid out from shot L-2 to M-5 on the wide-angle line. Reflection lines carried out from 1994 to 1997 in the southernmost part of the HCZ and refraction/wide-angle reflection lines in 1984 and 1992 are also shown. SYB: Sorachi-Yezo Belt; KMB: Kamuikotan Metamorphic Belt; IB: Idon'nappu Belt; HMB: Hidaka Metamorphic Belt; HB: Hidaka Belt; YB: Yubetsu Belt; TB: Tokoro Belt; HMT: Hidaka Main Trust. The Sorachi-Yezo Belt (SYB) is divided into the 110-100 Ma by the protrusion of serpentinite melange Sorachi Group, the Yezo Super Group and the including high-P metamorphic rocks. Kamuikotan Metamorphic Belt (KMB), whose for- Many thrusts have an almost N-S trend in the mation is directly related to the late Jurassic to early SYB, IB and the western part of the HB (Fig. 2). Cretaceous subduction process. The upper part of the These regions are under a high compressional stress Sorachi Group consists of early Cretaceous volcano- state resulting from two collision regimes (Kimura et genic sandstone, basaltic pillow lava, acidic tuff and al., 1983; Kimura and Miyashita, 1986). One is the siliceous shale (Kito et al., 1986; Kito, 1987). The dextral oblique collision between the North American Yezo Super Group, which conformably overlies the (Okhotsk) and Eurasian Plates associated with Oligo- Sorachi Group, is mainly composed of flysch type cene-middle Miocene back-arc spreading, by which a sedimentary rocks, shale and conglomerate, reflecting new eastward subduction of the Eurasian Plate the depositional environment of an early to late occurred along the central axis of Hokkaido Island. Cretaceous forearc basin. The KMB is characterized This subduction regime, whose existence is supported by serpentinites and high-P type metamorphic rocks by many geological and geophysical observations (Nakagawa and Toda, 1987). According to Nida and (e.g. Den and Hotta, 1973; Kimura et al., 1983), is Kito (1986), the prototype of this belt was created at considered to have continued up to the Pliocene and T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 63 then jumped to the eastern margin of the Japan Sea at migration of the KA and its collision with the NJA 1-2 Ma (Nakamura, 1983). Along this boundary, two (Kimura et al., 1983; Kimura, 1986). This arc-arc large earthquakes (M>8) occurred in the last 20 years. collision is considered to be responsible for the uplift From earthquake mechanisms and horizontal defor- of the metamorphic rocks forming the HMB. mation rates in this region, however, Sato and Ikeda (1994) pointed out that the eastward subduction within Hokkaido still continues to the present. The 3. Data acquisition and characteristics of seismic highly deformed folds and thrusts observed along the records western edge of the SYB were probably formed by a buried active fault system representing the plate The 1998-2000 Hokkaido Transect was aimed at boundary associated with this subduction. The second clarifying the crustal evolution and deformation collisional regime is due to the oblique subduction of processes associated with the ongoing arc-arc colli- the Pacific Plate under the Kuril Trench since the middle Miocene, which caused the southwestward tion of seismic refraction/reflection data and natural HidakaMts 6.0 5.0 4.0 1.0 -220.0 -200.0 009-008-0001-0020-0001-0091-0081- -40.0 -20.0 0.0 DISTANCE(km) b West East 6.0 5.0 0.0 1.0 90.0 -80.0 -70.0 -60.0 -50.0 -40.0 -30.0 -20.0 -10.0 0.0 10.0 DISTANCEINKM Fig. 3. Examples of record sections. A reduction velocity of 6.0 km/s was employed. (a) Record section from shot L-6 showing severe first both sides of the shot. At offsets of 7-10 km, travel-time jumps of 0.5-0.7 s are evident. 64 T. Iwasaki et al. / Tectonophysics 388 (2004) 5973 earthquake observations (Iwasaki et al.,2000a,b, processed together with the reflection data to obtain 2003a,b;Katsumata et al.,2002;Murai et al.,2003) a more reliable crustal image under the fold-and- The seismic refraction/wide-angle reflection experi- thrust belt. The dense seismic array for earthquake ments were carried out in 1999 and 2000 to obtain an observations,which consisted of 46 temporary entire crustal model from the hinterland to the stations and 40 permanent stations belonging to foreland fold-and-thrust belt crossing the HCZ (Iwa- universities of Hokkaido,Hirosaki and Tohoku and saki et al.,2000a,2003a;RGES,2002a,b;Fig.2).The Japan Meteorological Agency, was operated from profile line in 1999 was 227 km in length, along which July 1999 to July 2001 using a satellite data transfer 297 digital recording systems observed six dynamite system(Katsumataetal.,2002,2003).Off the shots (L-1 to L-4, M-5 and L-6) with 100700 kg Pacific coast, 27 ocean bottom seismometers were charges. In the vicinity of each shot point, five to six also deployed in both 1999 and 2000 to investigate additional receivers were deployed with a nearly 100 the detailedseismic activity and3-D velocity m spacing to measure a surface velocity. The western structure by covering a wider range of the HCZ half of this profile was surveyed again in 2000 to (Muraiet al.,2003). determine the detailed crustal structure under the fold Fig.3 shows examples of observed sections. First and-thrust belt. On the 2000 line,327 receivers were arrivals from shot L-6 (Fig. 3a) are characterized by deployed to record four explosive shots (L-1A and S-2 severe travel-time undulation of 0.20.5 s.Rather to S-4) with a charge size of 100300 kg. prominent phases following the first arrivals are wide- The seismic reflection profiling，which was angle reflections from within the crust, whose high undertaken from 1998 to 2000,was intended to apparent velocities indicate that the corresponding map a detailed crustal deformation caused by the reflectors are steeply inclined eastward.The record collision (Iwasaki et al.,2000a,b,2001b).The section from S-3 fired in the fold-and-thrust (Fig. 3b) profilesrecordedover 3years totaled138km in shows significant travel-time delays on both the sides length,and were laid out in the central part of the of the shot point, indicating the existence of a thick 1999 wide-angle line (between L-2 and M-5, see Fig. sedimentary layer of 35 km.We also notice travel- 2). As seismic sources, we used both dynamite and time jumps (abrupt delays) of 0.50.7 s at offsets of vibroseis shots.The refraction/wide-angle reflection 712 km. Such travel-time jumps are also found in data in the easterm half of the 2000 line were also record sections from shots L-2 to L-4 and S-2 to S-4, West NE Japan Arc Kuril Are East → Fold-and-Thrust Belt HMT HCZ Hinterland L-1A2.5-2.7S-2L-2 S-3 V5.9 1.6-1.7M-S 2.5-2.7 L-6 5.0-5.6 3.2-3.7 5.6 5.8 4.2-4.5 5.9 5.5 5.0.5.2 (y) 5.8 4504.8-53 6.3 6.0 5.9 6.1 0'9 Depth 10 5.3-5.6 6.3-6.5 5.8 6.3 6.2 1: 6.2 6.3-6.4 20 口 Middle/Lower Crust? Lower Crust 6.3 Reflective ower Crust 30 20 40 60 80 100 120 140 160 081 200 220 Distance (km) Fig. 4. Upper and middle crustal structure model across Hokkaido from the refraction/wide-angle reflection data. Note eastward dipping reflectors at a depth of 1025 km beneath the hinterland and velocity reversals within a thick sedimentary package developed in the fold-and- thrust belt. Beneath this sedimentary package, the crust of the NJA is subducted from the west. Open circles indicate hypocenters relocated within a 10 km distance from our profile by the dense earthquake monitoring array of the *Hokkaido Transect". T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 65 suggesting the wide occurrence of velocity reversals tionmethod such as travel-time inversion because of 50 the computational instability arising from the high Travel-times of the first arrival and prominent later degree of structural complexity. In the present phases for all of the shots were carefully picked and analysis, we did forward modelling based on a 2D graded with respect to their qualities (RGES, ray-tracing technique (Iwasaki, 1988), in which the 2002a,b). As shown in Fig. 3, the recorded travel- 50 timeshave a very complicated behaviour including ing head waves and diffracted waves from the edges undulations and significant jumps. This situation of layer boundaries as well as conventional diving and makes it difficult to apply a sophisticated interpreta- reflected waves. Uncertainties of model parameters a 5.0L-6 West East 四 4.0 3.0 RI 6.0 2.0 T 1.0 Rank A(<0.01s) RankB(<0.03s) 0.0%.0 Rank C(<0.05s) 20.0 40.0 60.0 80.0 100.0 0 64 220.0 Distance (km) b West East L-6 0.0 5.0 10.0 15.0 20.0 25.0 30.0 0.0 20.0 40.0 60.0 80.0 100.0 120.0 140.0 160.0 180.0 200.0 Distance (km) West East 4.0 M-5 (se 2.0 1.0 Rank A(<0.01s) RankB（<0. Rank C(<0.05s) 0.0L 0.0 20.0 40.0 60.0 80.0 100.0 120.0140.0 160.0 180.0 200.0 220.0 3 reading errors ranked by a grey scale. Solid lines indicate travel-times calculated from our model (Fig. 4). Phases R1 and R2 are wide-angle reflections from steeply dipping interfaces, whose ray paths are shown in (b). (b) Ray diagrams for shot L-6. (c) Travel-time plots of shot M-5. 50 66 T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 were estimated by examining data fit for velocity reflectors steeply dipping eastward. The geometry values and layer thicknesses slightly different from the of thesereflectors wasmainlyconstrained from final model. travel-time data of phases R1 and R2 from shots L- 6, M-5 and L-4 (Figs. 5 and 6). As shown in Fig. ite 4. Crustal model Fig. 4 shows our crustal model. The travel-time the fold-and-thrust belt where the geological struc- 1 V explained by highly deformed sedimentary layers with P-wave velocities of 1.6-2.7 and 3.5-3.6 km/s. model), the velocity abruptly increases and a high Their total thickness amounts to 44.5 km. At a velocity body (>6 km/s) crops out with horizontal depth of 10-25 km east of the HCZ, we notice extent of 10-20 km. Uncertainties of velocity a West East 11.0 L-4 10. 8. 7. 00 T-D/6. 5. 3.0 2.0 1.0 0. 0.0 10.0 20.0 30.0 40.0 Distance(km) 6 East L-4 0 25.0 30.0 150.0 160.0 170.0 180.0 190.0 200.0 Distance (km) Fig. 6. Record section and ray diagrams for shot L-4. (a) Record section with calculated travel-times (solid lines). A reduction velocity of 6.0 km/s is employed. Phases R1-2 and R5-6 are reflections from the delaminated KA crust, while R3-4 are those from the NJA crust as shown in (b). (b) Ray diagrams. T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 67 values are 0.05-0.15 km/s for the shallower part package includes one or two velocity reversals (Fig. (<10 km) and 0.15-0.2 km/s for the deeper part 4). Generally, the velocity and thickness of a low (10-25 km). The estimation error for the reflector velocity layer cannot be uniquely determined from depth is 2-3 km. The obtained structure provides travel-time data alone. In the present analysis, we seismological evidence for the westward obduction assumed its velocity value to be slightly smaller (3.2 of middle/lower crustal materials of the KA crust or 4.5 km/s) than that of the overlying layer, and corresponding to the western part of the HB. At a roughly determined the layer geometry from the depth of nearly 25 km, flat and westward dipping observed travel-times of both the first and later reflectors are also found forming a wedge-like arrivals. Figs. 7 and 8 show travel-time comparisons structure. A shot gather of the seismic reflection for shots S-3 and L-2, respectively. Later phases (R7 data from shot L-4 (Fig. 7) shows several reflec- and R8) within an offset range of 20 km were tions (R7-R10) with a different apparent velocity, interpreted as reflections from the bottom of the low which are mapped to interfaces with various dips at velocity layers. In Fig. 8, we also show a travel-time depths of 18-27 km (see also Fig. 4). curve of the diffracted wave from an edge of the The velocity model under the fold-and-thrust belt sedimentary package. The travel-times observed in is characterized by a thick (5-8 km) sedimentary the fold-and-thrust belt were explained well by our package with a velocity of 2.5-4.8 km/s. As modelling. The uncertainties in velocity are 0.1-0.2 indicated in the section from S-3 (Fig. 3b), this km/s in the shallower part (<5 km) while 0.2-0.3 km/s at depths of 5-15 km. The depth of layer interfaces may have an ambiguity of 2 km. a West East The upper 15-km part of the crust at the western 5.0 S-3 end of the profile is characterized by layers with a 4.0 R10 relatively high velocity (5.0-5.6 and 5.8-6.2 km/s), R9 probably corresponding to the NJA crust. Their estimation errors are about 0.1-0.2 km/s. These /6.0 layers are being subducted eastward beneath the R8 2.0 thick sedimentary package. This structure was mainly determined from the data of shots L-1, L-2 1.0 and L-1A (e.g. phases R10-12). Fig. 9 shows a Rank A(<0.01s) RankB(<0.03s） comparison between observed and calculated travel- times for shot L-1, in which the predominant phases 60.0 70.0 80.0 90.0 100.0 110.0 were interpreted as diving and reflected waves within Distance(km) the subducted NJA. The eastward dipping reflector b West East R12 within the NJA crust is also observed for shot S-3 L-2 (Fig. 8). Furthermore, phases appearing at offsets of 80-120 km from this shot are explained fairly 0.0 well as refracted and reflected waves within the NJA (km) 5.0 crust. 10.0 5. Discussion 20.0 60.0 70.0 80.0 90.0 100.0 110.0 The significant lateral structural changes in our Distance (km) crustal model are attributed to the ongoing arc-arc Fig. 7. Travel-times and ray diagrams for shot S-3. (a) Travel-time collision in this region. Fig. 10 shows a schematic plots with a reduction velocity of 6.0 km/s. Phases R7 and R10 are interpretation for our model, where geological infor- interpreted as reflections from the thick sedimentary package in the mation provided by Ito (2000), Arita et al. (2001) and fold-and-thrust belt and the subducted NJA crust (see also (b)). (b) Ray diagrams showing interpreted reflection phases R7 to R10. Kazuka et al. (2002) is also included. 68 T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 5. 1.0 0.0 60.0 -40.0 -20.0 0.0 20.0 40.0 60.0 80.0 100.0 120.0 140.0 Distance (km) East 5.0 10.0 15.0 20.0 25.0 30.0 0.0 20.0 40.0 60.0 80.0 100.0 120.0 140.0 160.0 180.0 200.0 220.0 Distance (km) Fig. 8. Record section and ray diagrams for shot L-2. (a) Record section with calculated travel-times (solid lines). A reduction velocity of 6.0 km/s is employed. Phase identifications for R7 to R12 are the same as given in Fig. 7(b). (b) Ray diagrams showing interpreted reflection phases R7 to R12. The HMB is considered to represent the middle) corresponds to a non-metamorphic sedimentary layer lower crustal part of the KA thrust up westward by the (the Nakanogawa Group), while the 5-7 km thick collision (Kimura, 1986, 1990; Kimura and Miya- 5.5-5.9 km/s layer represents meta-sediments occur- shita, 1986). Based on the metamorphic rocks in the ring at depths of 4-11 km in the petrological model. HMB, Komatsu et al. (1986) reconstructed a petro- The rocks from 11- to 17-km depth are considered to logical model for the KA crust prior to obduction. An be gneisses heavily intruded by tonalite-granodiorite upper 23 km of crust in their model consists of a non- and gabbro-diorite. The mixture of these rocks metamorphic sedimentary layer (the Nakanogawa explains our velocity of 6.1-6.4 km/s in almost the Group) and four metamorphic sequences ranging same depth range (Christensen, 1982; Christensen and from greenschist to granulite facies. On the 1992 Mooney, 1995). seismic line (Fig. 2), the structure beneath the HB The steep eastward-dipping reflectors east of the shows good correspondence to this petrological model HCZ give direct evidence for the obduction of the KA. (Iwasaki et al., 1998). In our seismic model, the upper These reflectors are also clearly imaged in the seismic 15-17 km crust east of the HCZ (at a distance range of reflection data (Adachi, 2002; Iwasaki et al., 2003a,b). 160-180 km in Fig. 4) is composed of 1.6-2.7, 3.5- On their westward extension, a rather high-velocity 3.6, 5.0-5.2, 5.5-5.9 and 6.1-6.4 km/s layers. The body (>6 km/s) crops out east of the HMT. Its uppermost two layers are interpreted as the Neogene horizontal extent is limited (<10-20 km), probably sedimentary rocks. The 5.0-5.2 km/s layer probably corresponding to the western part of the HB. The P- T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 1.0 0.0 0.0 20.0 40.0 60.0 80.0 100.0 120.0 140.0 160.0 180.0 200.0 220.0 Distance (km) b West East L-1 5.0 10.0 15. 20.0 [R12) 0.0 20.0 40.0 60.0 80.0 100.0 120.0 140.0 160.0 180.0 200.0 220.0 Distance (km) Fig. 9. Record section and ray diagrams for shot L-1. (a) Record section with calculated travel-times (solid lines). A reduction velocity of 6.0 wavevelocitydecreasesabruptlyto5.5-5.6km/s 1998 near-vertical reflection data indicate a high Vp/ eastward towards the hinterland. Furthermore, the Vs value (>1.80) in this body (Iwasaki et al., 1998; 1992refraction/wide-anglereflectiondata and the Adachi, 2002). These results strongly support the East West NE JapanArc Kuril Arc Fold-and-ThrustBelt HMT HCZ Hinterland L-1 S-2L-2 HidakaMts. L-1A S-3 S-4L-3 L-4 M-5 L-6 (y) NJA Crust KA Crust 10 Upper Crust Depth reflector Upper/MiddleCrust(BrittlePart) 20 Middle/LowerCrust? Lower Crust cactivity ？ LowerCrust ilePart?) 30 0 20 40 60 80 100 120 140 160 180 200 220 Distance (km) Fig. 10. Geological interpretation of the seismic model. KMB: Kamuikotan Metamorphic Belt; IB: Idon'nappu Belt; HMB: Hidaka Metamorphic Belt; Yz: Yezo Super Group; Sr: Sorachi Group; HMT: Hidaka Main Thrust. 70 T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 westward obduction model and exposure of relatively detachments (Kazuka et al., 2002). The seismic concentration of shallow microearthquakes (Fig. 10). activity under the fold-and-thrust belt (Katsumata with upper crustal rocks. et al., 2002) is quite low (Fig. 4). This implies that The flat and westward dipping reflectors at a depth is perhaps caused by asseismic slip on the deep of about 25 km under the hinterland (at 140-160 km package has progressed in an anelastic manner along in Fig. 4) are also mapped from the reflection data the above-mentioned detachments. (Adachi, 2002; Iwasaki et al., 2003a,b). Such reflec- Based on the above interpretation, we estimate the tion patterns imply that the crust of the KA is Range where the upwelling flow in the mantle wedge delaminated into more than two segments beneath 100 km across the fold-and-thrust belt. If the collision the study area. The delamination style imaged shows a started at 15-20 Ma, the shortening rate becomes 5-6 significant difference from that in the southernmost mm/year, which amounts to 60-70% of the total part of the HCZ (Arita et al., 1998; Ito, 2000, 2002; convergence rate (8 mm/year, Seno et al. (1996)) Fig. 2), where the KA crust is clearly separated into between the North American and Eurasian (Amurian) two segments. The dense passive seismic array of the Plates. Although this estimate of crustal shortening "Hokkaido Transect’ relocated hypocenters around a ing imaging possible at higher spatial resolution along the Backbone Range including the upper crust checked against other geophysical and geological which could not be understood from the image data, a large rate of 3-4 mm/year is also reported upper crust, producing a shallow inland earthquake from seismic reflection data in the southern part of the KA. These features indicate that the upper 25 km of HCZ (Ito, 2002, see Fig. 2) and in the fold-and-thrust crust (the obducted part) of the KA retains brittle belt (Kazuka et al., 2002). Hence, it can be said that properties. According to the seismic reflection data, the arc-arc collision in this area is still playing an causes numerous shallow microearthquakes as it Northern Miyagi earthquake (M6.5). Vp, Vs and Vp/ Range as predicted by the present model. Local two plates at the present time. Fig. 9 shows that, there is one more long, narrow The structure at the western end of our profile (Fig. brittle-ductile transition zone. 4) is similar to that of Northern Honshu. According to The most prominent structural features beneath the the Naruko volcano, there is another low-velocity Vs ratios in east-west vertical cross-section along a and middle crustal parts of Northern Honshu are 5.8-- eastward subduction of the NJA crust. Travel-time 6.25 and 6.2-6.45 km/s, respectively, and their total reaches the Moho. This region, in northern Miyagi thickness is 15 km. Furthermore, a reflective crust is along the previous seismic lines conducted in 1984 beneath the Backbone Range to immediately below and 1992 (Iwasaki et al., 1998, Fig. 2), indicating the also see a strong reflector at the bottom of the 6.0-6.2 velocity reversal occurs in a wide area of the fold-and- km/s layer. Therefore, the eastward dipping portion is thrust belt. The direct support for our interpretation is interpreted to be a subducted part of the NJA crust. provided from vertical seismic profiling (VSP) data 30 km north of our profile, which shows that an early Paleogene sedimentary unit (the Ishikari Group) 6. Conclusions appears three times in a 4.47-km deep well (Japan National Oil Corperation, 1999; Kazuka et al., 2002). A multidisciplinary project along the “Hokkaido The P-wave velocities of the deeper two layers are A.Hasegawa et al. / Tectonophysics 403 (2005) 59-75 4.06 and 3.43 km/s, forming velocity reversals at the determine structural inhomogeneities associated with depths of 2.8 and 3.8 km, respectively. From the good the arc-arc collision ongoing in Hidaka region, correspondence between this VSP information and Central Hokkaido, Japan. The collision of the Kuril our velocity structure, we interpret the low velocity Arc (KA) with the NE Japan Arc (NJA) since the layers under our profile line to be the same geological zone that branches off and extends to the eastern side unit of the Ishikari Group. The overlapping structures the Hidaka Mountains and the westward obduction of of the Ishikari Group were probably formed by the the Hidaka Metamorphic Belt (HMB). An E-W process of collision, where certain layer boundaries refraction/wide-angle reflection line from the KA to within the thick sedimentary package have acted as in the upper crust immediately above. Anelastic T. Iwasaki et al. / Tectonophysics 388 (2004) 59-73 71 and 2000, provided important constraints of crustal then stress will be concentrated in the area between deformation processes dominating the collision. disciplinary project, the “Hokkaido Transect". They In the eastern part of the profile, the KA crust is also thank to Prof. N. Hirata, Earthquake Research covered with 44.5 km thick highly deformed Neo- Institute, the University of Tokyo, and participants in gene sedimentary layers, beneath which two eastward the 1999 seismic reflection survey of this project for dipping reflectors are imaged in a depth range of 10-- their agreement to our usage of their records in the 25 km. Furthermore, the HMB is situated on the present work. The authors are indebted to Prof. H. western extension of these reflectors with physical Sato, Earthquake Research Institute, the University of properties consistent with middle/lower crustal rocks Tokyo, Prof. T. Ito, Faculty of Science, Chiba (a relatively high velocity (>6.0 km/s) and high Vp/Vs University and Dr. T. Kazuka, Dokkyo Saitama High (>1.80)). These results provide seismological support School for their valuable suggestions and comments for the westward obduction of middle/lower crust of during this study. The authors also express their the KA. Beneath the eastward dipping reflectors, thanks to Dr. K.I Katsumata, Graduate School of deeper flat and westward dipping reflector sequences Science, Hokkaido University, for kindly providing are situated at 25 and 25-27 km depths, respectively. the results of the natural earthquake observations The interpreted layer geometry probably indicates that mation in this region as well as in the Backbone the KA crust is delaminated into more than two Valuable suggestions and comments from Dr. aftershocks of the 1962 Northern Miyagi earthquake Randy Keller and another anonymous reviewer are distribution obtained by Katsumata et al., (2002), the appreciated. obducted part of the KA crust is considered to The experiments were undertaken by the Funds of maintain brittle properties. The seismic reflection the Special Works of the Earthquake Research data, on the other hand, show that the descending Institute, the Earthquake Research Institute, the Range but also to the hypocentral region of the 1962 University of Tokyo, as one of the disciplines of the 2003a,b). These indicate that the crustal delamination this area, the lower boundary of the seismogenic layer is occurring at the brittle-ductile transition zone of the KA crust. The western part of the profile across the NJA, References which belongs to the fold-and-thrust belt of the collision zone, is characterized by a very thick (5-8 Adachi, K., 2002. 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Murai, Y., Akiyama, S., Katsumata, K., Takanami, T., Yamashina, Research Group for Explosion Seismology, 2002b. Seismic the validity of this model must await future verifica- refraction/wide-angle reflection experiment across the foreland Shimamura, H., Furuya, I, Zhao, D., Sanda, R., 2003. area of the Hidaka collision zone, Hokkaido (Ohtaki-Biratori McKenzie, D.P., 1969. Speculations on the consequences and profile). Bull. Earthq. Res. Inst., Univ. Tokyo 77, 173-198 (in Urakawa-oki earthquake. Geophys. Res. Let. 30, 43-1-43-4. Japanese with English abstract). (10.1029/2002GL016459). Sakakibara, M., Niida, K., Toda, H., Kito, N., Kimura, G., Tajika, J., Nakagawa, M., Toda, H., 1987. Geology and petrology of Yubari- Kato, T., Yoshida, A., and the Research Group of the Tokoro dake serpentinite melange in the Kamuikotan tectonic belt, Belt, 1986. Nature and tectonic history of the Tokoro belt. Matsuzawa, T., Igarashi, T., Hasegawa, A., 2002. Characteristic Monogr. Assoc. Geol. Collab. Jpn. 31, 173-187. Japanese with English abstract). Sato, H., Ikeda, Y., 1994. Where is a boundary between Eurasian Nakamura, K., 1983. Possible nascent trench along the eastern and north American plates in northern Japan? 1994 Annual Kono, T., Nida, K., Matsumoto, S., Horiuchi, S., Okada, T., Matsumoto, S., Hasegawa, A., 1996. Distinct S-wave reflector in North American Plates. Bull. Earthq. Res. Inst. 58, 721-732 (in Programs 26 (7), p. A208. Japanese with English abstract). Seno, T., Sakurai, T., Stein, S., 1996. Can the Okhotsk plate be Nida, K., Kito, N., 1986. Cretaceous arc-trench systems in Davies, J.H., Stevenson, D.J., 1992. Physical model of source lio, Y., 1998. ohmin—a role on the generation of earthquake. J. 101, 11,305-11,315. (in Japanese with English abstract). Tsumura, N., Ikawa, H., Ikawa, T, Shinohara, M., 1999. Delami- Osanai, Y., Miyashita, S., Arita, K., Bamba, M., 1986. 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Journal of Mineralogical and Petrological Sciences,Volume 115, page 313-321, 2020 Geochemistry and magmatic zircon U-Pb dating of amphibolite 0.703827 0.703374 Yuji ICHIYAMA*, Takahito KOSHIBA*f, Hisatoshi ITO** and Akihiro TAMURA*** *Graduate School of Sciences, Chiba University, Chiba 263-8522, Japan **Civil Engineering Research Laboratory, Central Research Institute of Electric Power Industry, Abiko 270-1194, Japan Present address: Nittetsu Mining Consultants Co., Lid. Tokyo 108-0014, Japan Early Paleozoic serpentinite melanges in Japan preserve the oldest high-P metamorphic rocks in the circum- Pacific orogenic belt. To understand the tectonic regime at the subduction initiation of the proto-Japan con- phibolite blocks in the Omi serpentinite mélange, central Japan. The studied amphibolites from two different localities have the mineral assemblage of albite + clinozoisite + amphibole ± rutile ± titanite, which characterize epidote-amphibolite facies metamorphism. Whole-rock trace element concentrations of the amphibolites sug- gest that gabbroic protoliths formed possibly in an oceanic setting. The zircon U-Pb weighted mean ages ages of the studied amphibolites are comparable with those of reported Early Paleozoic ophiolite and high- pressure rocks in Paleozoic serpentinite mélanges in Japan. This fact implies that the young hot oceanic crust was subducting into the East Asian convergent plate margin during the Cambrian. Keywords: Serpentinite melange, Amphibolite, Zircon U-Pb dating, Cambrian INTRODUCTION and regarded them as the oldest high-P metamorphic rocks in the circum-Pacific orogenic belt. The petrology 0.703447 and geochronology of the oldest metamorphic rocks in Ja- orogenic belt, where the continuous accretion of oceanic pan will provide crucial clues for understanding the sub- materials since Early Paleozoic has created a nappe pile, duction initiation and tectonic environment in the Early where valuable information of ancient subduction dynam- Paleozoic East Asian continental margin. ics is preserved. For example, Late Paleozoic high-P Recent zircon geochronological studies have deter- blueschist metamorphic rocks (350-280 Ma Renge meta- mined the possible protolith ages of Early Paleozoic met morphic rocks) underthrusts below an Early Paleozoic amorphic and ophiolitic rocks in Japan (Osanai et al. ophiolitic complex (Oeyama ophiolite), and they represent 2014; Kimura and Hayasaka, 2019). Serpentinite mé- a subducted oceanic plate and a Paleozoic supra-subduc- lange enclosing the tectonic blocks of various lithologies tion zone complex, respectively (Ishiwatari et al., 2003) is distributed in the Omi area, northern central Japan (Fig. In addition to the Late Paleozoic high-P metamorphic 1a). In the Omi serpentinite mélange, the U-Pb ages of rocks, older 480-400 Ma amphibolites with high-P assem- blage are closely associated with Early Paleozoic serpen- 0.703954 tinite mélanges (Tsujimori, 2010). Tsujimori (2010) refer- za and Goto, 2010; Kunugiza et al., 2017). However, the 0.703778 geochronological information of amphibolite blocks in the Omi serpentinite mélange has been poorly character- doi:10.2465/jmps.191205 Y. Ichiyama, ichiyamay @ earth.s.chiba-u.ac.jp Corresponding au- ized so far. In this study, we performed whole-rock geo. 1ou 0.703843 314 Y. Ichiyama, T. Koshiba, H. Ito and A. Tamura Omi MS20-03 Japan 5km Cretaceous to Paleogene volcanics 500m and Neogene to Quaternary sediments Mushikawa Complex Sediment cover Serpentinite Cretaceous sediments Metagabbro Jurassic sediments Himekawa/Omi Complex Permian sediments Serpentinite melange Basic and pelitic schist Figure 1. (a) Geotectonic map of the Omi area (after Nagamori et al., 2010). The rectangle indicates the area shown in (b). (b) Geological map along the Omi River (after Matsumoto et al., 2011) and sample locations of this study. olites to infer the tectonic environment in the Paleozoic occurrence of glaucophane-bearing eclogite and blues- East Asian continental margins. chist in hosted in paragonite-bearing pelitic schists and the Non-eclogitic Unit with dominant biotite-bearing pel- GEOLOGICAL OUTLINE itic schists (e.g., Tsujimori, 2002). Recent boron isotope study for the serpentinite found a difference in isotope The Omi area comprises Paleozoic accretionary com- signature among the two units (Yamada et al., 2019). plexes of a broad age range; the Mushikawa, Hime- The Omi serpentinites contain antigorite, tremolite, talc, kawa-Omi, and Omi serpentinite mélange complexes carbonate, and metamorphic olivine (Yokoyama, 1985). from east to west in younging order (Fig. la; Nagamori In the Kotaki and Happo-One areas nearby the Omi area, et al., 2010). These complexes are equivalent to the Mai- the equivalent serpentinites locally experienced contact zuru (Permian), Akiyoshi (Carboniferous), and Oeyama metamorphism (Nozaka, 2003; Machi and Ishiwatari (Early Paleozoic) + Renge (Late Paleozoic) belts in the 2010). The protoliths of these serpentinites are harzburgite Inner Zone of southwest Japan, respectively. These Paleo- to dunite with rare chromitites (Yokoyama, 1985; Tsuji- zoic complexes are covered with Jurassic and Cretaceous mori, 2004; Machi and Ishiwatari, 2010). Machi and Ishi- sediments. The Omi serpentinite mélange is a tectonic watari (2010) suggested that these ultramafic rocks are melange composed mainly of mafic and pelitic schists similar to those reported from the Oeyama ophiolite and and amphibolite blocks with a serpentinite matrix (Fig. are derived from the residual mantle in Early Paleozoic 1b). Basic and pelitic schists experienced high-P meta- forearc regions. The studied amphibolites occur as several morphism (Banno, 1958; Shinji and Tsujimori, 2019; blocks of 100's of meters in dimension in the serpentinite Shinji et al., 2019), and eclogite facies metabasite (e.g.. matrix. Hornblende in the amphibolites yielded the K-Ar Tsujimori, 2002) and jadeitite (e.g., Tsujimori and Har- ages of 336 ± 13 and 370 ± 12 Ma (Shibata, 1981). low, 2017) also occur as tectonic blocks. The K-Ar ages of 339-285 Ma are recorded in phengitic muscovites from SAMPLE DESCRIPTION pelitic schists (Kunugiza et al., 2004). The Late Paleozoic HP schists are divided into the Eclogitic Unit with the In this study, we collected samples from two large am- Geochemistry and magmatic zircon U-Pb dating of the Omi amphibolites 315 b a Figure 2. Photographs of polished hand specimens for (a) MS20-01 and (b) MS20-03, and photom icrographs under plane-polarized light for (c) MS20-01 and (d) MS20-03. Rulers in (a) and (b) have 1 cm 0.5 mm. Ab, albite; Czo, clinozoisite; Amp, amphibole; Rt, rutile; Tnt, titanite. phibolite blocks from the Omi serpentinite mélange (Fig. MS20-01 at a given Mg# range (0.86-0.94). These min- 1b). These blocks are marked as *metagabbro' in the geo- eral assemblages indicate that they have experienced epi- logical maps provided by the previous studies (Banno, dote-amphibolite facies metamorphism. 1958; Nakamizu et al., 1989; Matsumoto et al., 2011; Nagamori et al., 2010). Both samples (MS20-01 and ANALYTICAL METHODS MS20-03) are coarse-grained and exhibit strong gneissic fabric, which is defined by the foliation of prismatic pla- We determined whole-rock major and trace elements of gioclase and amphibole and their compositional banding the studied amphibolites using an X-ray fuorescence (Figs. 2a and 2b). MS20-01 (from Kanayamadani) con- spectrometer (XRF; Rigaku ZSX Prims II) at Earthquake sists of green-color amphibole, clinozoisite, and albite Institute, The University of Tokyo. Glass beads made with a minor amount of magnetite and titanite (Fig. 2c). from the mixture of 1.000 g rock powder and 5.000 g The amphibole crystals in MS20-01 are classified as mag- flux (Li2B4O7) were prepared for the analysis. The de- nesiohornblende to tremolite-actinolite based on its Si = tailed analytical procedures and methods followed Hoka- 7.30-7.79 p.f.u. (per formula unit calculated on the basis nishi et al. (2015). of O = 23) and Mg#[=Mg/(Mg + Fe2+)] of 0.87-0.93. Whole-rock trace element, including rare-earth MS20-03 (from Shimizukura) consists of green-color am- element (REE), was measured using a laser-ablation in- phibole, clinozoisite, and albite with a minor amount of magnetite and rutile (Fig. 2d). Prehnite and albite veins MS) at Kanazawa University. A 193 nm ArF excimer are common in both samples. Albite is partially altered. laser (GeoLas Q+) and quadrupole mass spectrometer The amphibole crystals in MS20-03 are also composition- (Agilent7500s) were employed in this analysis. Direct- ally classified as magnesiohornblende to tremolite-actino- fused glasses made from rock powder were prepared lite but have lower Si (6.67-7.60 p.f.u.) than those of for the analysis. The detailed analytical procedures and 316 Y. Ichiyama, T. Koshiba, H. Ito and A. Tamura Table 1. Results of the whole-rock analyses using XRF and The detailed analytical procedures and methods followed ICP-MS those of Ito (2014). Note that we employed a laser spot XRF analysis ICP-MS analysis size of 20 μm with ~ 7 J/cm2 fluence, and the zircon standard 91500 (Wiedenbeck et al., 1995) was adopted MS20-1 MS20-3 MS20-1MS20-3 as a primary reference zircon instead of the Fish Canyon (wt%) (ppm) Tuff (Ito, 2014). We performed duplicate analysis (with SiO2 49.13 47.68 Li 0.83 3.54 TiO2 30 μm spot size) for the zircon grains from MS20-03 to 0.24 0.42 B 6.11 6.46 ensure the reliability of our measurement. The average Al2O3 13.75 15.71 Sc 40.71 39.93 reproducibility fo 206bb/238U age was 96%. The data ob- FeO* 4.86 5.92 Ti 1637 3333 tained in the second analysis, which show lower discord- MnO 0.11 0.12 V 170.0 208.2 ance% {= [(207Pb/235U age)/(206pb/238U age) - 1] × 100} MgO 11.28 10.80 Cr 835.8 274.6 than those of the first one, are preferentially used for dis- CaO 14.73 12.22 Co 39.70 39.42 cussion. A calculation software, Isoplot 4.0 (Ludwing, Na2O 1.81 2.25 Ni 211.4 127.1 2003), was used for processing data and preparing dia- K2O 0.12 0.23 Rb 1.79 1.38 grams. The U-Pb ages of the Plesovice zircon was meas- P2O5 0.03 0.01 Sr 192.5 127.7 ured as a secondary standard during the ICP-MS analy- Total 96.06 95.36 Y 10.19 12.28 ses. The weighted mean 206Pb/238U ages of the Plesovice Zr 9.35 18.85 zircon obtained through the two analyses were 336.6 ± (udd) Nb 0.24 0.58 6.3 (MSWD = 2.1) and 338.6 ± 2.8 (MSWD = 0.9), Sc Cs 0.51 0.14 36.3 32.7 which were in concordance with the recommended value V 167.8 197.4 Ba 28.41 48.26 Cr 1004.6 433.4 0.43 (337.13 ± 0.37 Ma; Slama et al., 2008). The analytical La 0.57 Co 34.8 34.6 Ce 1.56 2.04 results of the samples and Plesovice zircon are listed in Ni 240.4 137.0 Pr 0.30 0.38 Supplementary Tables S1 and S2 (available online from Cu 39.1 76.1 PN 1.89 2.32 https://doi.org/10.2465/jmps.191205), respectively. Zn 25.0 31.7 Sm 0.84 0.99 The cathodoluminescence (CL) images of the pol- Ga 10.0 12.5 Eu 0.41 0.53 Rb 0.2 IPq Gd 1.31 ished zircon grains were obtained using a CL detector 1.53 Y 10.7 12.9 Tb 0.25 0.29 (Oxford MiniCL) equipped with a scanning electron mi- Sr 181.2 121.6 Dy 1.81 2.11 croscope (JEOL SEM JSM-5600) at Chiba University. Zr 5.2 9.0 Ho 0.38 0.44 Nb IPq IPq Er 1.15 1.37 Ba RESULTS 31.1 46.0 Tm 0.16 0.19 Pb 2.4 0.6 Yb 1.07 1.32 Th IPq IPq Lu 0.21 0.21 Whole-rock geochemistry Hf 0.32 0.56 Ta 0.03 0.06 There is no significant difference in major element com- Th 0.02 0.01 position between the two amphibolite samples. SiO2 and FeO*, total iron calculated as FeO. MgO contents recalculated to anhydrous are slightly less bdl, below detection limit. than 50 and 11 wt%, respectively. FeO*/MgO ratios are low (<1.0). Chondrite-normalized REE patterns show de- pletion in light REE and positive anomalies in Eu (Fig. methods followed Tamura et al. (2015). The results of the 3a). Mid-ocean ridge basalt (MORB)-normalized trace whole-rock analyses using XRF and LA-ICP-MS are element patterns display distinct positive anomalies in listed in Table 1. Rb, Ba, Pb, and Sr (Fig. 3b), which are highly mobile Zircon grains for U-Pb dating were recovered by during secondary chemical modification. All MORB- hand picking after separations using heavy liquid (sodium normalized values, except for these anomalously enriched poly-tungstate) and neodymium magnet. The recovered elements are lower than 1. Negative anomalies in Zr and zircon grains were mounted into PFA fluorocarbon poly- Hf are present in both samples. The pattern of MS20-1 mers sheets and were polished with diamond paste. also exhibits slight negative anomalies in Nb and Ti. Zircon U-Pb dating was carried out using an LA- ICP-MS at Central Research Institute of Electric Power Zircon CL image and U-Pb dating Industry. A 213 nm Nd-YAG laser (NewWave Research) and sector-field mass spectrometer (Thermo Fisher Sci- We performed the U-Pb dating and CL observation of 18 entific ELEMENT XR) were employed in this analysis. and 14 grains for MS20-01 and MS20-03, respectively. Geochemistry and magmatic zircon U-Pb dating of the Omi amphibolites 317 have thin overgrowth rims (Fig. 4). (a) MS20-01 In general, concordance between the 207pb/235U and MS20-03 206pb/238U ratios obtained from the studied amphibolites MAR gabbro is not sufficient enough for a meaningful discussion (Figs. 5a and 5b), and, particularly, the results from MS20-01 Ch are lacking reliability. This comes from the high errors of urement and also possible common Pb. Only two analyses Y from MS20-01 and 11 from MS20-03 show discordance of less than 10%. The weighted mean 206Pb/238U ages cal- LaCePrNdSmEuGdTbDyHoErTmYbLu culated from the concordant zircon grains are 573 ± 29 Ma (mean square weighted deviation (MSWD)< 0.1) and 483 (b) ± 17 Ma (MSWD = 4.6) for MS20-01 and MS20-03, re- spectively (Figs. 5c and 5d). Instead, when the discord- ance is increased to 20%, the weighted mean 206pb/238U ages result in 541 ± 28 Ma (7 grains; MSWD = 2.0) and 477 ± 13 Ma (17 grains; MSWD = 4.3) (Figs. 5c and 5d), which are consistent with the above ages within the error ranges. In this study, the 206Pb/238U age of483 ± 17 Ma is considered for MS20-03 as a reliable geochronological data, and the 206Pb/238U age of 541 ± 28 Ma for MS20- 01 is treated as reference information. 001RbaThNbLaCePr rNmz HFEu TiGTby HErTmYbLu DISCUSSION Figure 3. (a) Chondrite-normalized REE patterns of the Omi am- phibolites (b) MORB-normalized trace-element patterns of the The whole-rock geochemistry of the studied amphibo- Omi amphibolites. Normalization values of chondrite and MORB are from McDonough and Sun (1995) and Hofmann lites is characterized by the depletion in trace element (1988), respectively. content. The REE patterns are similar to that of MORB in terms of light REE depletion (Fig. 3a). The Eu positive anomalies were probably caused by plagioclase accumu- The observed zircon grains generally show stubby and lation, indicating that the protolith of these amphibolites anhedral shape, and some of them have fuid inclusions is gabbroic rocks. On the other hand, the Zr and Hf neg- and irregular cracks. The zircon grains from MS20-01 do ative anomalies are detected in the trace element patterns not show any distinct zoning patterns under CL observa- (Fig. 3b). In particular, MS20-1 also has slight Nb and Ti tion. The CL images of the zircon grains from MS20-03 negative anomalies, which imply supra-subduction zone exhibit oscillatory and sector zonings, and some of them magmatism (e.g., Pearce and Peate, 1995). However, (a) (b) MS20-03-01 MS20-03-25 Figure 4. Representative cathodoluminescence images of the analyzed zircon grains from MS20-03. Scale bars are 40 μm. 318 Y. Ichiyama, T. Koshiba, H. Ito and A. Tamura 0.12 (a) b) 0.11 206Pb/238U 0.08 0.09 0.08 0.07 0.07 MS20-01 MS20-03 1.4 0.06 0.6 1.0 1.8 2.2 2.6 0.4 0.6 0.8 1.0 1.2 207Pb/235U 207Pb/235U 640 640 (C) (d) 600 (Ma） Discordance% <±20 Discordance% <±10 Ma 477±13Ma 483± 17 Ma (MSWD = 4.3) (MSWD = 4.6) age age 520 D 52 573 ± 29 Ma 48 (MSWD <0.1) Discordance%<±20 440 541±28Ma (MSWD = 2.0) MS20-01 MS20-03 400 400 () e I-o () po se ue pim Nz/dooz pe co-o (q) pe Io-o (e) go sod io SqnssM- s a MS20-03. such Nb, Zr, and Ti negative anomalies are also observed ing epidote-amphibolite facies metamorphism. The ob- in gabbroic rocks recovered from fracture zones in the tained Cambrian U-Pb ages indicate the ages of zircon Mid-Atlantic Ridge (Fig. 3b; e.g., Godard et al., 2009). crystallization from magma, i.e., the solidification ages These anomalies are attributed to plagioclase accumula- of the protolith of the studied amphibolites. It is uncertain tion because of the high incompatibility of these elements what the age gap of more than 50 m.y. between the two against plagioclase (Drouin et al., 2009). In addition, zir- samples means at this moment. con fractionation could have also been involved in Zr and Tsujimori (2010) regarded 480-400 Ma amphibo- Hf anomalies. Although it is difficult to identify which lites associated with Early Paleozoic serpentinite mélang- tectonic settings are suitable for the protolith of the stud- es in Japan as the oldest rocks that experienced high-P ied amphibolites, a spreading axis in an oceanic basin or metamorphism. The amphibolites are generally character- a backarc basin would be more likely. ized by epidote-amphibolite facies, but the Fuko Pass The zircon U-Pb data of the studied amphibolites amphibolites from Kyoto Prefecture, southwest Japan, yield Cambrian ages. The CL images of the zircon grains bear kyanite, paragonite, and rutile, and their protolith analyzed exhibit zoning indicating crystallization in an is interpreted as olivine and plagioclase cumulate (Tsuji- igneous environment (Fig. 4) but not the characteristics mori and Ishiwatari, 2002; Tsujimori and Liou, 2004). of metamorphic and inherited zircons. The zircon Th/U The presence of Al-rich clinopyroxene and spinel pseu- ratios ranging from 0.4 to 1.2 support that they crystal- domorph of corundum-magnetite symplectite in the Fuko lized from igneous processes but not metamorphic events Pass amphibolites implies that the rocks have experi- that are expected to be as low as <0.1 (e.g., Rubatto, 2017). Zircon in high-grade metamorphic rocks, such as amphibolite facies, and the subduction of the thick ocean- granulite and eclogite, exhibit homogeneous metamorphic ic crust was inferred (Tsujimori and Liou, 2004). The overgrowth (Corfu et al., 2003). The thin overgrowth rims 443-403 Ma hornblende K-Ar ages of the Fuko Pass am- observed in the studied zircon grains possibly formed dur- phibolites were interpreted as the ages of a subduction- Geochemistry and magmatic zircon U-Pb dating of the Omi amphibolites 319 related metamorphic event (Tsujimori et al., 2000). The in the Cambrian. The protolith ages of the studied am- studied amphibolites lack exact high-P mineral assem- phibolites are comparable with those of the Early Paleo- blages, but the assemblage of clinozoisite and rutile is zoic Oeyama ophiolite and the protolith of high-P rocks similar to that of the Fuko Pass amphibolites. The studied Sunon u ued u su adnas sozod u amphibolites have strong deformation texture and are dis- hot oceanic crust would have subducted in the East Asian tinct from non-subduction ophiolitic epidote-amphibo- convergent plate margin during the Cambrian. lites. For instance, epidote-amphibolites in the Late Perm- ian Yakuno ophiolite, southwest Japan, are generally lack- ACKNOWLEDGMENTS ing deformed texture and partially preserve primary igne- ous textures (Ishiwatari, 1985; Ichiyama and Ishiwatari, This study is part of the master thesis of T. K. at Chiba 2004). The studied amphibolites are possibly equivalent University. We are grateful to M. Tsukui and N. Furu- to the Fuko Pass amphibolites. kawa for their constructive comments. We thank N. Comparable age data have been reported from sev- Hokanishi for helping XRF analysis. This study was sup- eral localities in the Paleozoic formations in Japan. Tsuji- ported by JSPS KAKENHI Grant Number JP16K17831 mori et al. (2005) reported the zircon U-Pb ages of jade- and Itoigawa Geopark. The manuscript was improved by itite from the Osayama serpentinite mélange and inter- critical comments from T. Tsujimori and an anonymous preted them as a hydrothermal origin. Fu et al. (2010) reviewer. We also thank M. Satish-Kumar for his edito- investigated the oxygen isotopic zircon composition of rial handling and valuable comments. the Osayama jadeitite and identified igneous cores sur- rounded by hydrothermal rims in single zircon crystals. SUPPLEMENTARYMATERIALS The U-Pb ages of igneous cores and hydrothermal rims are not distinguishable in the range of 530-450 Ma, im- Supplementary Tables S1-S2 are available online from plying the subduction of young hot oceanic crust (Tsuji- https://doi.org/10.2465/jmps.191205. mori, 2010). The wide age variation of the hydrothermal zircon shows a correlation with initial epsilon hafnium REFERENCES values, suggesting the timing of a zircon source differen- tiated from the mantle at ~ 570 Ma (Tsujimori, 2017). Banno, S. (1958) Glaucophane schists and associated rocks in the Kunugiza et al. (2017) also confirmed the older inherited ages of 560 ± 16 Ma than the hydrothermal stage of 519 nal of Geology and Geography, 29, 29-44. 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(2017) Early Paleozoic jadeites in Japan: An over- struction of the oldest Pacific-type high P/T type metamor- view. Journal of Mineralogical and Petrological Science, 112, phism and hydration metamorphism during exhumation. Jour- 217-226. nal of Geography, 120, 4-29 (in Japanese with English ab- Tsujimori, T., Nishina, K., Ishiwatari, A. and Itaya, T. (2000) 403- stract). 443 Ma kyanite-bearing epidote amphibolite from the Fuko McDonough, W.F. and Sun, S. (1995) The composition of the Pass metacumulate in Oeyama, the Inner Zone of southwest- Earth. Chemical Geology, 120, 223-253. ern Japan. The Journal of the Geological Society of Japan, Nagamori, H., Takeuchi, M., Furukawa, R., Nakazawa, T. and 106, 646-649 (in Japanese with English abstract). Nakano, S. (2010) Geology of the Kotaki district. Quadrangle Tsujimori, T. and Ishiwatari, A. (2002) Granulite facies relics in the Series, 1:50,000. pp. 130, Geological Survey of Japan, AIST, early Paleozoic kyanite-bearing ultrabasic metacumulate in (in Japanese with English abstract). the Oeyama belt, the inner zone of southwestern Japan. Gond- Nakamizu, M., Okada, M., Yamazaki, T. and Komatsu, M. (1989) wana Research, 5, 823-835, Metamorphic rocks in the Omi-Renge serpentinite melange, Tsujimori, T. and Liou, J.G. (2004) Metamorphic evolution of ky- Hida Marginal Tectonic Belt, Central Japan. In High-pressure anite-staurolite-bearing epidote-amphibolite from the Early metamorphic bels and tectonics of the Inner Zone of southwest Palaeozoic Oeyama belt, SW Japan. Journal of Metamorphic Japan (Nishimura, Y., Hashimoto, M., Hara, I. and Watanabe, Geology, 22, 301-313. T. Eds.). The Memoirs of the Geological Society of Japan, 33, Tsujimori, T., Liou, J.G., Wooden, J. and Miyamoto, T. (2005) U- 21-35 (in Japanese with English abstract). Pb dating of large zircons in low-temperature jadeitite from Nozaka, T. (2003) Compositional heterogeneity of olivine in ther- the Osayama serpentinite melange, southwest Japan: insights Geochemistry and magmatic zircon U-Pb dating of the Omi amphibolites 321 into the timing of serpentinization. International Geology Re- ron isotope compositions of antigorite-grade serpentinites in view, 47, 1048-1057. the Itoigawa-Omi area of the Hida-Gaien Belt, Japan. Journal Tsujimori, T. and Harlow, G.E. (2017) Jadeitite (jadeite jade) from of Mineralogical and Petrological Sciences, 114, 290-295. Japan: History, characteristics, and perspectives. Journal of Yokoyama, K. (1985) Ultramafic rocks in the Hida marginal zone. Mineralogical and Petrological Science, 112, 184-196. Memoirs of the National Science Museum, 18, 5-18. Wiedenbeck, M., Allé, P., Corfu, F., Griffin, W.L., et al. (1995) Three natural zircon standards for U-Th-Pb, Lu-Hf, trace el- Manuscript received December 5, 2019 ement and REE analyses. Geostandards and Geoanalytical Manuscript accepted March 22, 2020 Research, 19, 1-23. Published online June 26, 2020 Yamada, C., Tsujimori, T., Chang, Q. and Kimura, J.-1. (2019) Bo- Manuscript handled by M. Satish-Kumar
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Ichiyama (2020) Geochemistry and magmatic U-Pb dating in Omi serpentinite melange.txt
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3 The Geology and Tectonics of Kyushu. Part 1: Tectonic Setting and Evolution from 150 Ma to 15 Ma Japan is part of the "Ring of Fire," the belt of earthquakes and volcanic activity that distinguishes the active margins of the Pacific ocean from the passive margins of the Atlantic ocean (Figure 3.1). In the 1920s, the great seismologist K. Wadati, discovered that a s o ba n n d u a sn near the Japan Trench to depths of about 5o0 km beneath the Japan Sea. This trench is a north-south trending bathymetric depression about 150 km east of the mainland that is as deep as 90o0 m. The towering volcanoes of northern Japan are centred about 75 km from one another and form a curving line that is about 1o0 km above the inclined seismic zone. In the 1930s, Japanese seismologists discovered that many of the earthquakes near the Japan the Ryukyu trench that is as deep as 5200 m and extends 2200 km from Kyushu to Taiwan. The cause of the active tectonic movements affecting the Earth was not well explained until the late 1960s when the realisation was made that the outer part of the Earth is divided into pieces, known as plates, about 100 km thick (Takeuchi et al., 1970). The plates consist of both the crust and cool uppermost mantle that has sufficient long-term strength that a is si dsas m n e n nsn r creating new ocean crust in the process of seafloor spreading at ocean ridges and come trenches, inclined seismic zones, and lines of explosive volcanoes. Plates slide past one another at transform boundaries. GPS measurements delineate where tectonic motions are underway (Sagiya et al., 2ooo). To a first order, the current tectonics of the Japanese islands can be explained by the interaction of four plates: Pacific, Philippine Sea, Eurasian and North American (Figure 3.1). The eastern ns ' e m si d o slightly different speed and direction than the parent plate (Wei and Seno, 1998; Heki et al., 1999). With the development of GPS geodesy in the 1990s, small differential subplate movements can be directly measured. The North American plate continues across the Bering Sea into eastern Asia and down past the Kamchatka-Kurile trench segments to Japan. An s o s d n o o o manifestation of the interactions between the Amur and Okhotsk subplates with the Pacific plate. Subduction along the Japan Trench at a speed of about 9 cm/a (90 km/Ma) is concurrent with convergence near the eastern edge of the Sea of Japan at a speed of 1 to 1.5 cm/a (10 to 15 km/Ma) (Okamura et al., 1995). The active tectonics of Kyushu at the northern end of the Ryukyu arc-trench system are the focus of this Case Study. The active tectonics of Southwest Japan (Kyushu, Shikoku, and southwest Honshu) are the manifestation of the o s convergence beneath Kyushu is at a speed of about 7 cm/a (70 km/Ma) (Seno et al., 1993; Zang et al., 2002). East-central Honshu, just south of Tokyo, near 34°N, is the convergent American (Okhotsk) plates. Prior to middle Cenozoic time, the entire length of the Japanese islands was underthrust by Pacific region and the Philippine Sea plate was created to the south. This plate is unique amongst the large plates in that it is nearly entirely surrounded by subduction zones. Because 27 of the lack of a spreading ridge connection to a surrounding plate, the movement history of the Philippine Sea plate is not well determined (Hall et al., 1995; Lee and Lawver, 1995). North American Eurasian plate Active plate Tectonic Boundary Amursubplate 140°E Okhotsk Strike-slip subplate <2cm/yr Japan Convergent Hokkaido > 2 cm/yr Sea Divergent Volcanoes Activesince1900AD 135°E Active 10,000 yrs BP Active Fault 40°N Wadati-BenioffZone Depth Contour inkilometers # 200 400 Japan kilometers Honshu S 35°N Shikoku Pacific 六 plate Kyushu Nankai Q Philippine Sea plate Shikoku Basin volcanoes that have been active in the last 10,0o0 years are also shown. Izu-Bonin-Mariana arc-trench system. The western edge of the Philippine plate subducts beneath Asia creating the Ryukyu arc-trench system. The Izu arc and triple junction present position. In so doing, subduction beneath Shikoku, and southwest Honshu changed the Nankai Trough. 28 3.1 Kyushu Basement Terranes Most of the basement of Kyushu is composed of four geologic terranes, from north to south: follows are mostly summaries from Kimura et al. (1991) and Taira (2001). The basement of the northern part of the island is primarily composed of the Sangun 1998). This terrane was intruded and widely metamorphosed by voluminous Jurassic to Tectonic Line. s s metamorphosed by voluminous Cretaceous (~100 to 80 Ma) arc plutons. The Higo subterrane, located only in west-central Kyushu, consists of greenschist to granulite facies rocks that were metamorphosed in the Triassic (Hamamoto et al., 1999). The extent of high temperature/low- pressure metamorphism varies greatly in intensity depending upon proximity to the Cretaceous plutons. The Ryoke belt is separated from the southern basement terranes by the Usuki-Yatsushiro Tectonic Line, an extension of the Median Tectonic Line that continues as a magmatic rocks occur south of the Usuki-Yatsushiro/Median Tectonic Line. In east-central Kyushu and south of the Usuki-Yatsushiro/Median Tectonic Line is a 10 km wide peninsula underlain by a wedge-shaped exposure of the Sanbagawa Belt. This outcrop area is the southwestern limit of the 7o0 km long belt that is well exposed in northern Shikoku s h ( o) g pressure/low-temperature blueschist facies metabasalts with lesser sediments recrystallised belts were recognised as the classic example of paired metamorphic belts (Miyashiro, 1961) that were later juxtaposed along a major strike-slip fault zone. The Sanbagawa Belt is in fault contact with the Chichibu Belt, a 10 to 20 km wide terrane that metamorphosed Carboniferous to Jurassic strata with local chaotic melange zones containing blocks of greenstone, chert, and limestones. The southern boundary of the Chichibu Belt is the Butsuzo Tectonic Line. The southern third of the island is composed of the Shimanto Belt, the late Cretaceous to the southern half of Shikoku and the Kii peninsula of Honshu (Taira et al., 1982; Mackenzie et al., 1987). The Shimanto Belt is composed of two main units: relatively coherent sequences of interbedded sandstone and shale and chaotically mixed units of shale-matrix melange containing slabs and blocks of basaltic lavas that are commonly pillowed, limestones and radiolarian cherts. 3.2 Tectonic History of Kyushu Region Most of the bedrock of the Japanese islands was assembled as a result of westward the late Paleozoic to early Jurassic that are mostly composed of deep ocean sediments by folding, faulting, and flowage during offscraping and underplating and then locally blanketed by slope deposits. In the Mid-Jurassic, between about 175 to 135 Ma, a major period of strike-slip faulting occurred along the margin of Asia that created the Oita- Kumamoto Tectonic Line and other high-angle fault contacts between displaced pieces of the late Paleozoic to early Jurassic accretionary prism. 29 In the very latest Jurassic or earliest Cretaceous, a major phase of fast subduction began that continues to this day off northeast Japan. This generated the enormous volume of arc magmas that invaded the Ryoke and Higo terranes and caused widespread high- s ssn oo si s overlying sediments that was imbricated and underplated beneath the leading edge of the new subduction zone and thoroughly metamorphosed and intensely deformed under high- pressure/low-temperature conditions. The Chichubu Belt and the older part of the Shimanto Belt are mostly trench sediments that were offscraped or slope deposits that accumulated at the same time. All of this accretion and metamorphism was caused by the subduction of the Pacific plate beneath the eastern edge of Asia. Between 50 to 40 Ma, a profound tectonic event occurred. The Pacific plate changed direction from northwest to a more westerly trend. This change is recorded as the major kink (dated at 43 Ma) in the Hawai-Emperor seamount chain as well as several other seamount chains in the Pacific plate farther to the south. The change in plate motion was coeval with the initiation of subduction along a major northwest-trending transform zone extending from New Zealand to a location north of the Equator. This formed the Izu-Bonin-Mariana and Tonga-Kermadec arc-trench systems (Hall et al., 1995; Lee and Lawver, 1995). Whether the initiation of d d n nns Izu peninsula that was established somewhere between Taiwan and Kyushu has migrated northwards ever since (Figure 1). The TTT triple junction appears to have reached the southern tip of Kyushu at about 25 Ma (Kimura et al., 2005). From about 32 to 23 Ma, Japan began to rift from Asia (Jolivet et al., 1994). Seafloor spreading in the Sea of Japan created new ocean crust in the backarc region between 23 to 12 Ma as Southwest Japan (southern Honshu and Shikoku) rotated clockwise about 45 degrees and Northeast Japan shifted eastward with a small counterclockwise rotation (Ishikawa, 1997). As this occurred, the TTT triple junction migrated more eastward than northward resulting in highly oblique and very slow convergence beneath southwest Japan. This movement caused significant left-lateral strike-slip offset along the Median Tectonic Line. As the TTT triple junction migrated northwards, the plate interactions along southern Japan abruptly switched from the subduction of the Pacific plate to the Philippine Seaplate. -i a moo i e o pa s Mariana arc split with backarc spreading creating the Shikoku Basin between 25 to 15 Ma (Karig, 1974; Sdrolias et al., 2004). The Kyushu-Palau Ridge is the remnant arc, nearly 400 km behind the Izu-Bonin-Mariana arc. The opening of the Shikoku Basin created new ocean crust that was still hot as it was underthrust beneath Kyushu and Shikoku islands and the Kii peninsula. The combination of spreading and underthrusting accounts for the minor but very anomalous forearc magmatism from 17 to 12 Ma (Hibbard and Karig, 1990; Kimura et al., 2005) and the widespread, low-temperature metamorphism of the Shimanto Belt on Kyushu 3.3Summary The overall geologic history that created the basement terranes of Kyushu is well understood as a result of the long-term westward subduction of the Pacific ocean seafloor. Subduction in the late Paleozoic to mid-Mesozoic created the Sangun accretionary prism, which makes up the northern part of the island. This accretionary event was followed by a period of major 30 strike-slip transform faulting between about 175 to 135 Ma that shuffled the accretionary prism and made some of the high angle fault zones that are mapped as tectonic lines. The reinitiation of westward subduction in the very latest Jurassic or earliest Cretaceous generated voluminous magmatism that invaded the northern half of the island and the large accretionary prism that makes up most of the southern half of Kyushu. The major plate motion change at about 43 Ma that isolated the Philippine Sea plate led to the northwards migration of the trench-trench-trench triple junction that is now located at the Isu peninsula. As the triple Ryoke belt with abundant Mesozoic plutons to become juxtaposed directly against the plate generated the Ryukyu trench and arc. 3.4 References for Section 3 s migrated seismic Sections: American Association of Petroleum Geologists Memoir 34, p. 309-322. Banno, S., and Nakajima, T., 1992, Metamorphic belts of the Japanese islands: Annual Reviews of Earth and Planetary Sciences, v. 20, p. 159-179. Chai, B. H. T., 1972, Structure and tectonic evolution of Taiwan: American Journal of Science, V. 272, p. 389-422. Cloos, M., Sapie, B., Quarles van Ufford, A., Weiland, R. J., Warren, P. Q, and McMahon, T. breakoff: Geological Society of America Special Paper 400, 51 pp. Cox, A., and Engebretson, D., 1985, Change in motion of Pacific plate at 5 Ma BP: Nature, v. 313, p. 472-474. Hall, R., Ali, J. R., Anderson C. D., and Baker, S. J., 1995, Origin and motion history of the Philippine Sea plate: Tectonophysics, v. 251, p. 229-250. Hamamoto, T., Osanai, Y., Kagami, H., 1999, Sm-Nd, Rb-Sr, and K-Ar geochronology of the Higo metamorphic terrane, west-central Kyushu, Japan: The Island Arc, v. 8, p. 323-334. Heki, K., Miyazaki, S., Takahashi, H., Kasahara, M., Kimata, F., Miura, S., Vasilenko, N., F., lvashchenko, and Ki-Dok, A., 1999, The Amurian plate motion and current plate kinematics in eastern Asia: Journal of Geophysical Research, v. 104, p. 29,147-29,155. Hibbard, J. P., and Karig, D. E., 1990, Structural and magmatic responses to spreading ridge Ishikawa, N., 1997, Differential rotations of north Kyushu Island related to middle Miocene clockwise rotation of SW Japan: Journal of Geophysical Research, v. 102, p. 17,729- 17,745. Jolivet, L, K., Tamaki, K., and Fournier, M., 1994, Japan Sea, opening history and mechanism: A synthesis: Journal of Geophysical Research, v. 99, p. 22,237-22,259. Kamata, H., 1989, Volcanic and structural history of the Hohi volcanic zone, central Kyushu, Japan: Bulletin of Volcanology, v. 51, p. 315-332. Kamata, H., and Kodama, K., 1994, Tectonics of an arc-arc junction: An example from Kyushu Island at the junction of the Southwest Japan Arc and the Ryukyu Arc: Tectonophysics, v. 233, p. 69-81. Karig, D. E., 1974, Evolution of arc systems in the western Pacific: Annual Reviews of Earth and Planetary Sciences, v. 2, p. 51-75. P n go ' i responses in SW Japan during Neogene time: Geological Society of America Bulletin, v. 117, p. 969-986. Kimura, M., 1985, Back-arc rifting in the Okinawa Trough: Marine and Petroleum Geology, v. 2, p.222-240. 31 o o s 287 pp. Kodaira, S., Takahashi, N., Park, J.-O., Mochizuki, K., Shinohara, M., and Kimura, S., 2000 s seismic survey: Journal of Geophysical Research, v. 105, p. 5,887-5,905. Kodama, K., and Nakayama, K.-l., 1993, Paleomagnetic evidence for post-late Miocene intra- arc rotation of south Kyushu: Tectonics, v. 12, p. 35-47. Lee, T.-Y., and Lawver, L. A., 1995, Cenozoic plate reconstruction of Southeast Asia: Tectonophysics, v. 251, p. 85-138. Mackenzie, J. S., Needham, D. T., and Agar, S. M., 1987, Progressive deformation in an accretionary complex: An example from the Shimanto belt of eastern Kyushu, southwest Japan: Geology, v. 15, p. 353-356. Malavieille, J., Lallemand, S. E., Dominguez, S., Deschamps, A., Lu, C.-Y., Liu, C.-S., Schurle, P., and the ACT Scientific Crew, Arc-continent collision in Taiwan: New marine 187-211. n 1 o o do 1 1 ' paleomagnetic evidence from the south Ryukyu Arc: Tectonophysics, v. 175, p. 335-347. Miyashiro, A., 1961, Evolution of metamorphic belts: Journal of Petrology, v. 2, p. 277-311. Nishimura, Y., 1998, Geotectonic subdivision and areal extent of the Sangun belt, inner zone of Southwest Japan: Journal of Metamorphic Geology, v. 16, p. 129-140. Okamura, Y., Watanabe, M., Morijiri, R., Satoh, M., 1995, Rifting and basin inversion in the eastern margin of the Japan Sea: Island Arc, v. 4, p. 166-181. Sagiya, T., and Thatcher, W., 1999, Coseismic slip resolution along a plate boundary 104, p. 1,111-1,129. Sagiya, T., Miyazaki, S., Tada, T., 2ooo, Continuous GPS array and present-day crustal deformation of Japan: Pure and Applied Geophysics, v. 157, p. 2303-2322. Sdrolias, M., Roest, W. R., and Muller, R. D., 2004, An expression of Philippine Sea plate s Eurasian plate: Tectonophysics, v. 42, p. 209-226. Seno, T., Stein, S., and Gripp, A. E., 1993, A model for the motion of the Philippine Sea plate 17,941-17,948. Sibuet, J.-C., Deffontaines, B., Hsu, S.-K., Thareau, N., Le Formal, J.-P., Liu, C.-S., and the ACT party, 1998, Okinawa trough backarc basin: Early tectonic and magmatic evolution: Journal of Geophysical Research, v. 103, p. 30,245-30,267. Sugimura, A., and Uyeda, S., 1973, Island Arc -- Japan and its Environs: Elsevier, Amsterdam, 247 pp. s Earth and Planetary Sciences, v. 29, p. 109-134. Taira, A., Okada, H., Whitaker, J. H., McD., and Smith, A. J., 1982, The Shimanto Belt of Japan: Cretaceous - lower Miocene active-margin sedimentation: Geological Society of London Special Publication 10, p. 5-26. Takeuchi, H., Uyeda, S., and Kanamori, H., 1970, Debate about the Earth: Approach to o a o s n so Francisco, 281 pp. 32 Utu, K., Hoang, N., and Matsui, K., 2004, Cenozoic lithospheric extension induced magmatism in Southwest Japan: Tectonophysics, v. 393, p. 281-299. Wallis, S., 1998, Exhuming the Sanbagawa metamorphic belt: The importance of tectonic discontinuities: Journal of Metamorphic Geology, v. 16, p. 83-95. Wei, D., and Seno, T., 1998, Determination of the Amurian plate motion: American Geophysical Union, Geodynamics series, v. 27, p. 337-346. Yanagi, T., Nakada, S., and Watanabe, K., 1992, Active volcanoes and geothermal systems in the fault zone of middle Kyushu: 2g" International Geological Congress (IGC) Field Trip A23, p. 1-31. Yamaji, A., 2003, Slab rollback suggested by latest Miocene to Pliocene forearc stress and migration of volcanic front in southern Kyushu, northern Ryukyu Arc: Tectonophysics, v., 634,p.9-24. Zang, S. X., Chen, Q. Y., Ning, J. Y., Shen, Z. K., and Liu, Y. G., 2002, Motion of the Philippine Sea plate consistent with the NUVEL-1A model: Geophysical Journal International, v. 150, p. 809-819. Zhao, D., Ochi, F., Hasegawa, A., Yamanto, A., 2ooo, Evidence for the location and cause of large crustal earthquakes in Japan: Journal of Geophysical Research, v. 105, p. 13,579- 13,594. 33
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tectonics_kyushu.txt
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An adakitic pluton on Kyushu Island, southwest Japan arc
Atsushi Kamei *
Research Center for Deep Geological Environments, Geological Survey of Japan, National Institute of Advanced Industrial Science and Technology,
Tsukuba Central 7, Tsukuba 305-8567, Japan
Received 20 December 2002; revised 24 June 2003; accepted 17 July 2003
Abstract
An adakitic pluton is designated the Shiraishino granodiorite ca. 15 km2in outcrop area on Kyushu Island, southwestern Japanese arc. Sr,
Y, and rare earth element concentrations and crystallization characteristics of the pluton are similar to those of adakites. Geochemical and
tectonic considerations show that the pluton was derived from partial melting of a subducted slab with minor contamination of high87Sr/86Sr
material and subsequent fractional crystallization of plagioclase þhornblende.
Compiled Sr and Y data indicate that most Cretaceous granitic rocks of southwestern Japan, generated by partial melting of mafic lower
crust. Adakitic plutons derived by slab melting occur rarely and have older ages. It is concluded that plutonism in the southwestern Japanesearc during Cretaceous time was began with the emplacement of small adakitic intrusions and progressed to more abandant granitic rocks ofthe type derived from the mafic lower crust.
q2003 Elsevier Ltd. All rights reserved.
Keywords: Adakite; Granodiorite; Southwest Japan; Geochemistry
1. Introduction
The southwestern Japanese arc contains numerous calc-
alkaline granitic plutons ( Fig. 1 ) (e.g. Takahashi, 1983 ). The
source of the granitic magma was mainly igneous rock, notsedimentary (e.g. Kagami et al., 1992; Nakajima, 1996 ).
Two possible candidates for the source are commonly
discussed: (1) a subducted slab (e.g. Maruyama et al., 1997 )
and (2) the mafic lower crust (e.g. Kagami et al., 1992 ).
In the case of slab melting, granitic magma with adakitic
composition is generated under pressures ( $2.2 GPa) high
enough to stabilize garnet and/or amphibole withoutplagioclase (e.g. Martin, 1987; Drummond et al., 1996 ).
Partial melting of mafic lower crust can also generateadakitic magma leaving a residue similar to that generatedin the case of slab melting under pressures of $1.6 GPa
(e.g. Atherton and Petford, 1993 ). At pressures ,1 GPa,
typical calc-alkaline granitic magma is produced leavingplagioclase þpyroxenes ^amphibole without garnet (e.g.
Johannes and Holtz, 1996 ). Adakites have higher Sr
concentrations and lower levels of Y compared with typicalcalc-alkaline granitic rocks because Sr and Y are highlycompatible elements for plagioclase and garnet, respect-
ively (e.g. Drummond et al., 1996; Martin, 1999 ). At depths
equivalent to 1–1.6 GPa, granitic magma between inter-mediate adakites and granites would be generated, becausethis magma is produced within the garnet þplagioclase
stability region (e.g. Rapp et al., 1991; Johannes and Holtz,
1996; Martin, 1999 ).
Most Cretaceous granitic rocks in southwest Japan have
low Sr and high Y contents when compared with adakitesbut adakitic plutons also exist. The examination of Sr andY compositions of the granitic rocks is key to identifyingtheir source.
A Cretaceous adakitic pluton has been recognized in
Kyushu, southwestern Japan. The aim of the present studyis (1) to determine its petrological characteristics, (2) todiscuss its petrogenesis, and (3) to consider the generalcharacteristics of Cretaceous granitic rocks in south-western Japan from the viewpoint of their Sr and Ycontents.
2. Cretaceous granitic rocks in southwest Japan
The Cretaceous granitic province of southwestern Japan
is divided into four units: Kyushu Island unit, Ryoke Belt,
1367-9120/$ - see front matter q2003 Elsevier Ltd. All rights reserved.
doi:10.1016/j.jseaes.2003.07.001
Journal of Asian Earth Sciences 24 (2004) 43–58
www.elsevier.com/locate/jaes
*Tel.: þ81-29-861-3787; fax: þ81-29-861-3643.
E-mail address: kamei-a@aist.go.jp (A. Kamei).
Fig. 1. Cretaceous granitic provinces in southwest Japan. Modified after Nozawa (1975) . The granitic province of San’in includes Paleogene–Neogene plutons.
A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 44
San’yo Belt, and San’in Belt ( Fig. 1 ). Granitic intrusions in
Kyushu took place between 121 and 76 Ma, and consist ofboth I- and S-types ( Osanai et al., 1993; Owada et al., 1999;
Kamei, 2002 ), mainly magnetite-series plutons ( Ishihara,
1981 ). Granitic rocks in the Ryoke Belt, ranging in age from
121 to 73 Ma, are characterized by gneissose texture(Kagami et al., 1992, 1999; Nakajima, 1996; Yuhara et al.,
2000 ). Granitic rocks in the San’yo Belt are less foliated and
range in age from 124 to 70 Ma ( Kagami et al., 1992, 1999;
Nakajima, 1996 ). Most of the Ryoke and San’yo Belts
granitic rocks are I-type ilmenite-series rocks ( Ishihara,
1981 ). On the other hand, granitic rocks in the San’in Belt
range from late Cretaceous to Neogene in age (85–20 Ma)and are mostly I-type magnetite-series ( Ishihara, 1981;
Kagami et al., 1992, 1999; Nakajima, 1996 ).
Recently, a province of small adakitic plutons ( #4k m
2)
was recognized in the central part of the San’yo Belt ( Kiji
et al., 2000 )(Fig. 1 ). These are older than the neighboring
granitic rocks in the San’yo and Ryoke Belts.
3. Cretaceous granitic rocks in Kyushu Island
Cretaceous granitic rocks are exposed in northern and
central Kyushu Island ( Fig. 2 ). They intrude various
lithologies, such as low- and high-pressure type meta-morphic rocks, accretionary complexes, and coevalvolcanic rocks. The southern limit of Cretaceous granitic
rock distribution is the Usuki–Yatsushiro tectonic line
(Fig. 2 ). Granitic rocks are widely exposed in the
north but are less prominent in the centralpart because of abundant Cenozoic volcanic cover(Sasada, 1987 ).
Cretaceous granitic rocks in Kyushu belong to two types:
(1) tonalite to granodiorite, and (2) granite ( Karakida, 1985;
Owada et al., 1999; Kamei, 2002 )(Fig. 2 ). These rocks
differ in their mineral assemblages; i.e. rocks of the tonalite
to granodiorite suite always contain hornblende, whereas
rocks of the granite suite contain mostly muscovite and/orgarnet with no hornblende. Tonalite to granodioritic rockswere emplaced before the granites ( Karakida, 1985; Owada
et al., 1999; Kamei, 2002 ). Recently, Kamei et al. (1999)
proposed that the tonalite to granodiorite magmas werederived from partial melting of mafic lower crust, and
Kamei (2002) inferred that the granite resulted from partial
melting of tonalitic middle crust.
The adakitic pluton outcrops at the southern end of a
group of Cretaceous granitic rocks, called the Shiraishino
granodiorite, and belongs to the Higo plutonic rocks ( Fig. 2 ).
The Shiraishino granodiorite belongs to the tonalite togranodiorite suite based on its petrographic composition(Kamei, 2002 ), even though its geochemical composition
appears different from those of other members of the
tonalite to granodiorite suite.
Fig. 2. Sketch map showing the distribution of Cretaceous plutonic rocks in Kyushu Island, southwest Japan. Box shows the location of Fig. 3 .A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 45
4. Local geology
Fig. 3 shows a geological map of the Higo plutonic rocks
and the surrounding Higo and Ryuhozan metamorphicrocks.
The Higo metamorphic rocks, mainly derived from
pelitic, psammitic, basic, and carbonate rocks, arecharacterized by high-grade metamorphism up to granulitefacies with anatexis ( Yamamoto, 1962; Obata et al., 1994;
Osanai et al., 1996 ). They reflect two important
metamorphic events at ca. 250 and ca. 100 Ma ( Osanai
et al., 1993, 1998; Nakajima et al., 1995; Nagakawa et al.,1997 ). The Ryuhozan metamorphic rocks experienced
low- to medium-grade metamorphism up to amphibolite
facies ( Yamamoto, 1962; Sakashima et al., 1999 ), and are
mainly derived from mafic and felsic tuff, carbonate rocksand minor clastic sediments. Early Permian fossils havebeen extracted from Ryuhozan carbonates ( Murata et al.,
1981 ). Clastic sediments have forearc basin or passive
margin characteristics and peak metamorphism is con-sidered as having taken place at ca. 100 Ma ( Sakashima
et al., 1999 ).
The Higo plutonic rocks consist of three bodies:
Shiraishino granodiorite, Manzaka tonalite, and Miyano-hara tonalite ( Yamamoto, 1962 )(Fig. 3 ). The Shiraishino
granodiorite is equigranular, whereas the Manzaka tonaliteis characterized by E–W flow structure defined by maficminerals. The Miyanohara tonalite possesses a myloniticfoliation striking around E–W. The degree of myloniticdeformation changes gradually from protomylonite in thenorth to mylonitic gneiss in the south ( Kamei et al.,
2000 ). The Shiraishino granodiorite intrudes both the
Miyanohara tonalite and the Higo metamorphic rocks. TheMiyanohara tonalite intrudes Ryuhozan metamorphicrocks. The Manzaka tonalite intrudes Higo metamorphicrocks. Other intrusive boundaries for the plutonic rockscannot be determined in the field because of the overlyingthe Aso welded tuff ( Fig. 3 ). Ages of the Higo plutonic
rocks are listed as follows; Shiraishino granodiorite:121^14 Ma (Rb–Sr whole rock age: Kamei et al.,
1997 ), Manzaka tonalite: 111.7 ^1.8 Ma (U–Pb
SHRIMP age: Sakashima et al., 1998 ), Miyanohara
tonalite: 210.8 ^1.2 Ma (Sm–Nd internal isochron age:
Kamei et al., 2000 ).
Owada et al. (1998) observed that Mesozoic granitic
plutons in Kyushu can be grouped into the three agecategories of Indosinian (300–200 Ma), Early Yanshanian(190–140 Ma), and Late Yanshanian (140–70 Ma), definedfor southeastern granites in China ( Jahn et al., 1990 ). The
Shiraishino granodiorite and the Manzaka tonalite, togetherwith all Cretaceous granitic rocks in Kyushu, are categor-ized on Late Yanshanian (Cretaceous) plutons, whereas theMiyanohara tonalite is an Indosinian pluton ( Owada et al.,
1998; Kamei et al., 2000 ).5. Petrography of the Shiraishino granodiorite
The Shiraishino granodiorite is medium- to fine-grained,
and mainly consists of plagioclase, quartz, biotite, horn-blende, and K-feldspar, with minor amounts of opaqueminerals, titanite, apatite, zircon, and rutile. Plagioclase andhornblende are euhedral to subhedral. Most of theplagioclase shows normal zoning (An ¼44–14). K-feld-
spar and biotite are euhedral to anhedral, whereas quartz isanhedral. Myrmekite commonly occurs, and chlorite,muscovite, and sericite are secondary minerals.
6. Analytical procedure
Modal compositions and major and trace element
compositions of the Shiraishino granodiorite are fromKamei et al. (1997) and are given in Table 1 . These
chemical data were analyzed using an X-ray fluorescence
(XRF) spectrometer. Based on repeated analyses ðn¼10Þ
of standards of Geological Survey of Japan, analytical errorswere estimated to be better than 3% for major elements, and10% for trace elements. Concentrations of some traceelements and rare earth elements (REE) ( Table 1 ) were
determined by inductively coupled plasma mass spec-trometry (ICP-MS; Activation Laboratories Ltd., Canada).The mineral compositions were determined using anelectron microprobe (Shimadzu V6) at the Center ofInstrumental Analysis of Yamaguchi University. In quan-titative analyses, an acceleration voltage of 15 kV, a
specimen current of 15 nA, and a spot beam width of
2mm were employed. The results of the mineral analyses
are presented in Table 2 .
7. Geochemistry
Based on compiled whole rock chemical compositions
from Cretaceous granitic rocks in southwestern Japan,variation diagrams and Sr/Y vs. Y diagram were constructed(Figs. 4 and 5 ). Plots of the Kyushu, Ryoke, San’yo, and
San’in regions occupy broad trends in the variationdiagrams ( Fig. 4 ). The plots of rocks in Kyushu are
subdivided into tonalite to granodiorite and granite. Thegranite in Kyushu has higher Al
2O3and Sr, and lower Y
compared with the tonalite to granodiorite rocks. Adakitesin the San’yo Belt, reported by Kiji et al. (2000) , are
enriched in Na
2O and Sr. On a Sr/Y vs. Y diagram, granitic
rocks of SW Japan mainly plot within the island arc
andesite–dacite–rhyolite (ADR) field, whereas some rocksof Ryoke Belt, adakites in the San’yo Belt, and the granite inKyushu plot within the adakite field ( Fig. 5 ).
In contrast to the Kyushu, Ryoke, San’yo, and San’in
regions, the Shiraishino granodiorite (SiO
2contents: 63.81–
68.20 wt%) is characterized by high Sr concentrations(439–730 ppm) and slightly higher Na
2O (3.67–4.10 wt%),A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 46
Fig. 3. Geological map of the Higo plutonic rocks in Kyushu Island.
A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 47
Table 1
Major and trace element, and modal data for the Shiraishino granodiorite
Sample No. 41510 42013 50602 50605 50702 50710 51007 51008 51009 51010 51011 51901A
SiO 2(wt%) 63.87a65.06a67.05a64.69a68.20a63.81a66.42a67.23a67.15a66.90a66.30a67.87a
TiO 2 0.66a0.63a0.49a0.57a0.44a0.66a0.49a0.47a0.49a0.56a0.62a0.44a
Al2O3 16.19a16.15a16.02a16.07a15.16a16.33a15.54a15.59a16.35a15.51a15.61a15.22a
Fe2O3 5.05a4.66a3.42a4.24a3.28a4.79a3.74a3.52a3.74a4.16a4.33a3.24a
MnO 0.10a0.09a0.07a0.08a0.07a0.09a0.09a0.07a0.08a0.09a0.09a0.07a
MgO 2.06a1.91a1.35a1.79a1.19a2.08a1.53a1.38a1.63a1.75a1.81a1.22a
CaO 4.86a4.49a3.76a3.62a3.35a4.83a4.04a3.78a3.86a3.01a3.44a3.50a
Na2O 3.96a3.67a4.10a4.02a3.87a3.92a3.92a3.97a3.71a3.80a3.73a3.91a
K2O 1.76a1.16a2.27a2.49a2.75a1.67a2.38a2.37a2.33a2.45a2.50a2.64a
P2O5 0.19a0.19a0.14a0.18a0.14a0.19a0.15a0.15a0.15a0.16a0.17a0.13a
LOI 1.64 1.78 0.97 1.65 0.68 1.09 0.98 0.83 1.18 1.49 1.71 1.70
Total 100.24 99.79 99.64 99.40 99.13 99.46 99.28 99.36 100.67 99.88 100.31 99.94
Ba (ppm) 422a261a315a709b480a515b248a571b312a416a456a587b
Cr 5a17a–a10a–a21a6a9a17a4a5a–a
Nb 17a12a11a14.0b11a12.7b12a13.3b11a11a13a12.2b
Ni 5a5a5a11a6a5a8a7a7a5a7a4a
Rb 45.3c38a61a63.2c77.4c38.4c63a62a55a56a62.3c77.4c
Sr 584c664a544a730c439c582c517a499a547a553a643c473c
V9 6a87a69a84a39a93a69a70a66a67a88a59a
Y2 0a10a16a13.3b18a12.7b20a13.0b15a17a16a10.4b
Zn 99a44a68a78a73a88a78a67a62a80a92a60a
Zr 181a218a168a201a176a170a162a169a169a162a187a159a
Hf 0.77b0.53b0.58b0.65b
Ta 0.83b0.71b0.89b0.79b
W 4.59b5.01b5.14b6.40b
Th 5.83b4.60b7.28b6.24b
U 1.19b0.95b1.34b1.17b
Mo 0.13b0.13b0.07b0.12b
La 23.9b23.3b28.1b30.3b
Ce 44.7b42.6b51.1b52.0b
Pr 4.50b4.22b4.99b4.77b
Nd 18.2b17.1b19.3b17.8b
Sm 3.37b3.06b3.26b2.84b
Eu 1.09b1.06b0.98b0.89b
Gd 2.81b2.61b2.69b2.30b
Tb 0.42b0.39b0.40b0.33b
Dy 2.10b2.00b1.96b1.62b
Ho 0.38b0.36b0.36b0.28b
Er 1.00b0.95b0.94b0.74b
Tm 0.15b0.14b0.14b0.11b
Yb 0.96b0.90b0.96b0.71b
Lu 0.16b0.14b0.15b0.11b
MODESQuartz (%) 24.2 26.0 28.0 25.6 33.1 25.9 23.6 32.9 24.3 22.6 28.2 30.0
Plagioclase 50.2 51.1 50.0 45.0 44.5 53.0 45.3 42.3 46.5 49.9 40.3 48.8K-feldspar 7.1 3.2 4.9 10.1 6.2 2.1 14.8 8.0 14.3 9.5 10.3 8.5
Hornblende 8.3 7.7 4.8 7.2 4.0 9.5 2.1 2.1 3.9 3.9 4.5 2.1
Biotite 9.9 11.0 10.3 10.0 11.0 7.2 13.3 13.3 9.7 12.7 14.0 10.0Opaque 0.2 0.5 1.5 1.5 0.5 2.1 0.5 1.0 0.5 0.5 0.5 0.3Titanite 0.1 0.5 0.5 0.6 0.7 0.2 0.4 0.4 0.8 0.9 2.2 0.3
Total iron as Fe
2O3. LOI is loss on ignition.b,cLOI and modal data are from Kamei et al. (1997) .
aAnalyzed by XRF Kamei et al., (1997) .
bAnalyzed by ICP-MS (new data).
cAnalyzed by isotope dilution method ( Kamei et al., 1997 ).A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 48
Table 2
Representative mineral compositions of the Shiraishino granodiorite
Sample No. 50710 Av. ðn¼8Þ51007 Av. ðn¼8Þ50710 Av. ðn¼8Þ51007 Av. ðn¼8Þ
Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim Core Rim
Plagioclase K-feldspar
SiO 2 59.47 61.12 57.86 58.60 59.00 58.91 61.53 59.51 61.70 59.80 65.08 65.38 64.99 64.80 64.98 64.81 64.99 64.85 64.63 64.79
TiO 2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.01 0.00
Al2O3 25.60 24.29 26.17 26.00 26.30 25.61 24.33 25.41 24.64 25.66 17.93 18.12 18.15 18.02 18.10 18.08 17.97 18.12 17.97 18.05
FeO 0.06 0.02 0.05 0.00 0.04 0.15 0.11 0.08 0.14 0.08 0.04 0.04 0.03 0.05 0.06 0.05 0.04 0.04 0.10 0.04
MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.03 0.01 0.00 0.01 0.00 0.00 0.00
MgO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.01 0.00 0.01 0.01 0.00 0.00 0.01 0.00 0.01CaO 7.79 6.34 8.94 8.62 7.85 8.66 5.90 7.07 5.86 7.05 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00Na
2O 7.18 8.03 6.39 6.66 7.07 6.00 8.39 7.40 8.39 7.76 1.16 1.26 0.85 0.81 0.90 1.06 1.09 0.98 0.91 0.99
K2O 0.10 0.20 0.05 0.07 0.05 0.10 0.11 0.18 0.20 0.06 15.64 15.74 15.86 15.89 15.69 15.86 15.84 15.54 15.67 15.83
Total 100.20 100.00 99.46 99.95 100.31 99.43 100.37 99.65 100.93 100.41 99.86 100.55 99.88 99.61 99.75 99.86 99.94 99.56 99.29 99.71
Hornblende Biotite
SiO 2 46.26 47.41 46.00 46.36 46.26 45.63 46.70 46.49 47.12 46.43 37.69 37.90 37.67 37.54 37.53 36.96 35.43 37.85 36.79 36.66
TiO 2 1.14 1.05 1.04 1.04 1.14 1.02 0.94 0.99 0.98 0.95 3.85 3.61 3.85 3.90 3.74 3.74 3.76 3.85 3.87 3.80
Al2O3 8.11 7.34 8.00 7.50 7.98 8.12 7.19 8.03 7.26 7.67 14.62 14.67 15.21 14.65 14.78 15.12 14.93 14.29 13.72 14.64
FeO 17.56 17.15 17.78 17.62 17.56 17.75 17.29 18.27 17.80 17.57 19.41 19.03 18.85 19.13 19.03 20.55 22.23 20.16 20.60 20.81
MnO 0.57 0.52 0.61 0.63 0.58 0.75 0.76 0.79 0.81 0.74 0.35 0.40 0.34 0.32 0.35 0.52 0.53 0.31 0.35 0.38
MgO 10.69 11.25 10.77 11.11 10.95 10.12 10.39 10.01 10.27 10.68 10.86 10.71 10.56 10.55 10.60 10.50 10.27 10.44 10.56 10.48CaO 11.22 11.34 10.83 10.80 11.01 10.85 11.03 10.96 11.21 11.14 0.02 0.00 0.07 0.09 0.02 0.07 0.05 0.02 0.00 0.04Na
2O 1.30 0.79 1.32 1.28 1.14 1.65 1.47 1.43 1.31 1.58 0.13 0.13 0.12 0.11 0.10 0.07 0.08 0.11 0.11 0.10
K2O 0.57 0.49 0.57 0.50 0.51 0.47 0.32 0.50 0.63 0.54 9.50 9.40 9.15 9.28 9.37 8.77 8.65 9.22 9.31 9.21
Total 97.42 97.34 96.92 96.84 97.13 96.36 96.09 97.47 97.39 97.30 96.43 95.851 95.82 95.57 95.52 96.30 95.93 96.25 95.31 96.12
Total iron as FeO.A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 49
and lower K 2O (1.16–2.64 wt%), Rb (38–77 ppm), and Y
(10–20 ppm) ( Fig. 4 ). InFig. 5 , the Shiraishino granodiorite
rocks plot within the adakite field. Chondrite normalized
REE patterns for the Shiraishino granodiorite show LREEenrichment and HREE depletion, and no marked Euanomaly [(La/Yb)
N¼16.9–28.8, Eu/Eu* ¼1.01–1.15]
(Fig. 6A ), resembling adakites and their Archaean equiva-
lents ( Fig. 6B ). The initial87Sr/86Sr (SrI) of the Shiraishino
granodiorite is 0.70493 ^7(Kamei et al., 1997 ).8. Discussion
8.1. Differentiation of the Shiraishino granodiorite
The Shiraishino granodiorite universally contains plagi-
oclase with normal zoning. Bulk SiO 2content increases
with decreasing average CaO content and increasingaverage Na
2O content of plagioclase ( Tables 1 and 2 ).
This suggests that differentiation of the Shiraishino
Fig. 4. Selected variation diagrams of major and trace element concentrations for Cretaceous granitic rocks in southwest Japan. Data are compiled fr omTakagi
et al. (1989), Takagi (1992), Yuhara (1994), Yuhara et al. (1998), Kamei et al. (1997), Kamei (2002), Yuge et al. (1998), Sugii and Sawada (1999), Kiji e t al.
(2000), Fujii et al. (2000), Ishihara and Wu (2001), Ishihara (2003a,b), and Kutsukake (2002) .A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 50
granodiorite resulted from fractional crystallization. The
proposed differentiation process was tested by using major
element mass balance modeling, subsequently, the model
was tested with trace element modeling.
The observed mineralogy was used for the major
element modeling. The minerals involved in fractionation
were determined on the basis of their relationship between
modal composition and bulk SiO 2content considered as a
differentiation index ( Fig. 7 ).Fig. 7 shows compositional
trends with inflections in the diagrams of K-feldspar,
biotite, and quartz. All points of the inflections aresituated at a SiO
2content of ca. 66 wt%. Consequently,
SiO 2#66 wt% and SiO 2.66 wt% will be referred to as
stage-1 and stage-2, respectively. During stage-1,plagioclase and hornblende decreased whereas biotiteand K-feldspar increased; quartz remained unchanged.
Consequently, during stage-1 only hornblende þ
plagioclase were extracted from the parent magma. Instage-2, modal compositions of plagioclase, hornblende,
K-feldspar, and biotite decreased indicating that they all
played a role in magmatic differentiation.
The results of major element modeling are given in
Table 3 . In stage-1, the compositions of initial and final
magmas were modeled using those of samples 50710(SiO
2¼63.81 wt%) and 51007 (SiO 2¼66.42 wt%),
respectively. The mineral compositions of the subtracted
phases in the modeling of stage-1 are the respective average
plagioclase and hornblende compositions of 50710(Table 2 ). Crystal fractionation of the assemblage plagio-
clase and hornblende in the weight proportions (59:41)
successfully predicted the composition of 51007 after
20.6% crystallization of 50710. The compositions of initialand final magma in stage-2 modeling were determinedbased on samples 51007 (SiO
2¼66.42 wt%) and 51901A
(SiO 2¼67.87 wt%), respectively. Mineral compositions
are average compositions of plagioclase, hornblende, K-feldspar, and biotite of 51007 ( Table 2 ). Crystal
fractionation of the assemblage plagioclase, hornblende,
K-feldspar, and biotite in weight proportions(58.3:40.6:0.2:0.9) successfully predicted the composition
of 51901A after 10.2% crystallization of 51007. The sum of
squares of residuals for calculations of stage-1 and stage-2does not exceed the acceptable upper limit ( ,1.5:
e.g. Mann, 1993; Tatsumi, 2001 ). In stage-2, although
K-feldspar (0.9%) and biotite (0.2%) are calculated, these
minerals cannot be regarded as important phases comparedwith those of plagioclase (58.3%) and hornblende (40.6%).
Therefore, it is concluded that the differentiation of
Shiraishino granodiorite mainly resulted from the fractionalcrystallization of plagioclase þhornblende through the
stage-1 and stage-2.
Trace element modeling based on the Rayleigh fractional
crystallization law was performed in order to confirm theresults of major element modeling. Results of major element
model were reintroduced in trace element modeling, and the
results are listed in Table 4 . Mineral proportions are
Fig. 5. Relationship between Sr/Y ratios and Y concentrations of Cretaceous granitic rocks in southwest Japan. Data sources are the same as Fig. 4 . Adakite and
Island Arc Andesite–Dacite–Rhyolite (ADR) fields are from Defant et al. (1991) .A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 51
Pl:Hbl ¼59:41 in both stage-1 and stage-2, and the degrees
of crystallization in stage-1 and stage-2 are 20.6 and 10.2%,respectively. Distribution coefficients used for models ofstage-1 and stage-2 were from tonalite to dacite and fromdacite to rhyolite, respectively ( Table 4 ). The starting
composition of trace element modeling for stage-1 was that
of 50710, and the results of stage-1 were input as the startingcomposition of stage-2 ( Table 4 ). The chondrite normalized
trace element pattern of 51901A was well reproduced by theresults of trace element modeling ( Fig. 8 ). Trace element
modeling supported the results of major element mass
balance modeling.8.2. Source of the Shiraishino granodiorite
Sr, Y, and REE concentrations of Shiraishino granodior-
ite resemble those of adakites ( Figs. 5 and 6A and B ). The
differentiation of granodiorite is mainly controlled by the
fractional crystallization of plagioclase þhornblende,
which is consistent with the crystallization system of
Archaean TTG described previously by Martin (1995) .
Fig. 6. Chondrite normalized REE patterns for the Shiraishino granodiorite
(A) and the adakites and their Archaean equivalents (B). The Shiraishinogranodiorite has similar REE patterns to the adakites. Data sources in (B)
areDrummond et al. (1996) (Adakite and TTD) and Martin (1995) (TTG
and Modern granite). Normalizing values are after Taylor and McLennan
(1985) .
Fig. 7. Variation diagrams of mineral modes vs. bulk SiO 2contents for the
Shiraishino granodiorite.A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 52
Adakitic magmas can be produced from partial melting
at the base of a thickened crust under conditions that
stabilize garnet and/or amphibole without plagioclase (e.g.
Atherton and Petford, 1993 ). This melting is generally
expected to occur at the high-pressure conditions of$1.6 GPa ( $48 km depth). The Shiraishino granodiorite
intruded into the forearc portion of Kyushu in EarlyCretaceous time, and many granitic rocks intruded in thebackarc side of the Kyushu after that. The area of south ofthe Usuki–Yatsushiro tectonic line of Kyushu ( Fig. 2 )
hardly existed when the Shiraishino granodiorite wasemplaced (121 ^14 Ma) ( Maruyama et al., 1997 ). The
present crustal thickness of Kyushu is estimated at ca.30 km based on seismic profiles ( Mitsunami, 1992 ). If the
crustal thickness of Kyushu has remained unchanged sincethe Early Cretaceous, adakitic magma could not have beenproduced from lower crustal melting, and it is unlikely thatthe crust was thicker in the Early Cretaceous than it is today.
I conclude that it is unlikely that the adakitic characteristics
of the Shiraishino granodiorite formed from the melting of a
thickened crust.
Adakites can be also produced by partial melting of
young ( ,25 Ma) subducted oceanic crust (e.g. Defant and
Drummond, 1990; Martin, 1999 ). Many authors have
suggested that the Kula-Pacific ridge migrated toward thesouthwestern Japanese arc in Early Cretaceous time ( Uyeda
and Miyashiro, 1974; Kinoshita and Ito, 1986; Nakajimaet al., 1990; Kinoshita, 1995; Maruyama et al., 1997 ). Based
onKinoshita (2001) , the ridge crest passed beneath the
present location of the Shiraishino granodiorite at ca.105.4 Ma, which is later than the igneous activity of thegranodiorite (121 ^14 Ma). This indicates that very young
slab was subducted at the time of granodioritic magmatism.Therefore, it is concluded that the Shiraishino granodioriteTable 3
Least squares fractional crystallization models using major elements for the Shiraishino granodiorite
Stage-1 Initial magma Final magma Subtracted phases
Plagioclase Hornblende
Sample No. 50710 51007 50710 50710
SiO 2(wt%) 63.81 66.42 59.00 46.26
TiO 2 0.66 0.49 0.00 1.14
Al2O3 16.33 15.54 26.30 7.98
FeO 4.31 3.36 0.04 17.56
MnO 0.09 0.09 0.00 0.58MgO 2.08 1.53 0.00 10.95
CaO 4.83 4.04 7.85 11.01
Na
2O 3.92 3.92 7.07 1.14
K2O 1.67 2.38 0.05 0.51
Calculated result 79.41% 12.14% 8.44%Degree of crystallization, 20.58%
Fractionated proportions 58.99% 41.01%
Sum of squares, 0.228
Stage-2 Initial magma Final magma Subtracted phases
Plagioclase Hornblende K-feldspar Biotite
Sample No. 51007 51901A 51007 51007 51007 51007
SiO
2(wt%) 66.42 67.87 59.80 46.43 64.79 36.66
TiO 2 0.49 0.44 0.00 0.95 0.00 3.80
Al2O3 15.54 15.22 25.66 7.67 18.05 14.64
FeO 3.36 2.92 0.08 17.57 0.04 20.81
MnO 0.09 0.07 0.00 0.74 0.00 0.38MgO 1.53 1.22 0.00 10.68 0.01 10.48CaO 4.04 3.50 7.05 11.14 0.00 0.04
Na
2O 3.92 3.91 7.76 1.58 0.99 0.10
K2O 2.38 2.64 0.06 0.54 15.83 9.21
Calculated result 89.75% 5.92% 4.13% 0.02% 0.09%Degree of crystallization, 10.16%
Fractionated proportions 58.26% 40.64% 0.22% 0.88%
Sum of squares, 0.018A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 53
Table 4
Rayleigh fractional crystallization models using trace elements for the Shiraishino granodiorite
Stage-1 Partition coefficient Initial magma (50710)
(ppm)Result of stage-1
(ppm)Sample 51007
(ppm)Stage 2 Partititon coefficient Initial magma (result
of stage-1) (ppm)Result of stage-2
(ppm)Sample 51901A
(ppm)
Plagioclase Hornblende Plagioclase Hornblende
Ba 0.160a0.090a515 629 248 Ba 0.360b0.044b629 684 587
K 0.110a0.330a13,862 16,671 19,756 K 0.263b0.081b16,671 18,192 21,915
Nb 0.025c1.300c12.7 14.1 12 Nb 0.060c4.000c14.1 13.1 12.2
La 0.130d0.200d23.3 28.3 – La 0.320e0.850e28.3 29.7 30.3
Ce 0.110d0.300d42.6 51.4 – Ce 0.240b0.899b51.4 54.2 52.0
Sr 2.000d0.360d582 539 517 Sr 2.840b0.022b539 501 473
Nd 0.070d0.800d17.1 19.8 – Nd 0.170b2.800b19.8 19.2 17.8
Sm 0.050d1.100d3.06 3.45 – Sm 0.130b3.990b3.45 3.19 2.84
Eu 1.300d1.300d1.06 0.99 – Eu 2.110b3.440b0.99 0.83 0.89
Gd 0.040d1.800d2.61 2.76 – Gd 0.900b5.480b2.76 2.28 2.30
Dy 0.031d2.000d2.00 2.08 – Dy 0.086b6.200b2.08 1.75 1.62
Y 0.600d1.900d12.7 12.3 20 Y 0.100c6.000c12.3 10.5 10.4
Er 0.026d1.900d0.95 1.00 – Er 0.084b5.940b1.00 0.85 0.74
aSource data of partition coefficients: Gill (1981) .
bSource data of partition coefficients: Arth (1976) .
cSource data of partition coefficients: Pearce and Norry (1979) .
dSource data of partition coefficients: Martin (1987) .
eSource data of partition coefficients: Henderson (1982) .A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 54
is derived from partial melting of young subducted oceanic
crust.
8.3. Petrogenesis of the Shiraishino granodiorite as
an adakitic pluton
Slab derived adakitic magma is generally rich in MgO,
Ni, Cr, Sr, and Na 2OþCaO ( Smithies, 2000; Martin and
Moyen, 2002 ), because melting occurs at greater depth
where plagioclase is not stable, and because of interactions
between the overlying mantle and magma ( Martin, 1999 ).
The Shiraishino granodiorite is poor in MgO(1.22–2.08 wt%), Ni (4–11 ppm), and Cr ( ,21 ppm)
compared with average adakite (MgO ¼2.20–1.96 wt%,
Ni¼24 ppm and Cr ¼36–46 ppm: Martin, 1999 ). Sr
(439–730 ppm) and Na
2OþCaO (6.81–8.82 wt%) con-
tents of the granodiorite are almost the same as that ofaverage adakite (280–706 ppm, 7.86–9.09 wt%: Martin,
1999 ). During fractional crystallization of plagioclase þ
hornblende, Mg, Ni, and Cr display compatible behavior
because the hornblende/liquid partition coefficients ðK
DÞare
high in felsic liquids ( Martin and Moyen, 2002 ). Although
Na, Ca, and Sr are compatible elements for plagioclase infelsic liquids, their K
Dvalues are not markedly higher than
those of hornblende ( Martin and Moyen, 2002 ). Conse-
quently, evolved adakite is expected to be quite poor inMgO, Ni, and Cr, and moderately depleted in Sr andNa
2OþCaO. Therefore, the geochemical features of the
Shiraishino granodiorite indicate that the granodiorite
crystallized from a slightly evolved magma affected by
fractional crystallization of plagioclase þhornblende.
The SrI of the Shiraishino granodiorite (0.70493 ^7) is
greater than that of mid-oceanic-ridge basalt (MORB)(0.7023–0.7034: Faure, 1986 ). Therefore, the granodioritemagma could not have been derived solely from partial
melting of subducted fresh oceanic crust. A previousisotopic examination for Shiraishino granodiorite indicatedthat the granodiorite is not affected by assimilation of high
87Sr/86Sr (0.7093–0.7125) wall rocks, and that the SrI of
granodiorite is controlled by its subcrustal source ( Kamei
et al., 1997 ). Therefore, the Shiraishino granodiorite has not
been influenced by an assimilation and fractional crystal-
lization (AFC) process with high87Sr/86Sr material during
solidification. The SrI of Shiraishino granodiorite could beexplained by the following processes: (1) the magmaformed by melting of a mixture of subducted slab andminor oceanic sediments, or (2) the magma was affected bylithospheric inputs before the start of crystallization.
The observations outlined indicate that the Shiraishino
granodiorite was derived from partial melting of subductedoceanic crust and was contaminated with minor high
87Sr/86Sr material. Subsequent fractional crystallization of
plagioclase þhornblende lead to an evolved magma.
8.4. Igneous activity of Cretaceous granitic rocks in
the southwestern Japanese arc
Cretaceous granitic rocks in southwestern Japan mostly
fall within the island arc ADR field in Sr/Y vs. Y diagram(Fig. 5 ). Sr contents of plutons in the San’in and San’yo
Belts are lower than 300 ppm ( Fig. 4 ), which does not
satisfy the definition of adakites of Martin (1999) . SrI of
plutons in the San’yo and Ryoke Belts is generally higherthan 0.7070, which closely resemble those of mafic rocks(gabbro, metadiabase, and mafic granulite) in the same
regions ( Kagami et al., 1992, 2000; Iizumi et al., 2000 ). This
indicates that the regional variation is atteributable tocompositionally different magma sources in the uppermantle or lower crust ( Kagami et al., 1992 ). Therefore, it
is concluded that most Cretaceous granitic rocks in south-west Japan were derived from partial melting of mafic lowercrust at pressure under ,1 GPa.
On the other hand, the majority of granites in Kyushu
plot within the adakite field ( Fig. 5 ), however, they contain a
lot of K-feldspar which is not common in adakites. Kamei
(2002) suggested that the granite magma was generated by
anatexis of metaluminous tonalite in the middle crust underconditions of variable fH
2O. Therefore, it is concluded that
they are not adakitic rocks formed by slab melting. Somegranitic rocks in Ryoke Belts also plot within the adakite
field ( Fig. 5 ), and might be generated at deeper levels that
stabilized garnet and/or amphibole.
The Shiraishino granodiorite, as an adakitic pluton, is the
oldest among Cretaceous granitic rocks on Kyushu. Kiji
et al. (2000) reported many adakitic bodies ( #4k m
2)
derived by slab melting in the central San’yo Belt. These arealso older than surrounding granitic rocks of the San’yo andRyoke Belts. Therefore, Cretaceous plutonism in south-western Japanese arc is considered to have started by theintrusion of small adakitic plutons that formed by slab
Fig. 8. Result of trace element modeling of fractional crystallization
compared with the observed values for the Shiraishino granodiorite.Normalizing values for chondrite are after Wood et al. (1979) and Taylor
and McLennan (1985) .A. Kamei / Journal of Asian Earth Sciences 24 (2004) 43–58 55
melting and terminated with the formation of larger granitic
rocks formed by the anatexis of the lower crust.
9. Conclusions
A 121^14 Ma adakitic pluton (Shiraishino granodior-
ite) is newly recognized among the Higo plutonic rocks ofKyushu in the southwestern Japanese arc. Sr, Y, and REEconcentrations of the pluton are similar to those of adakites.These characteristics indicate that the magma was generatedin the garnet and/or amphibole stability field. Geochemicalmodeling indicate, that differentiation of the pluton wasmainly controlled by fractional crystallization ofplagioclase þhornblende, which also corresponds to gen-
eral differentiation of adakites.
Based on tectonic considerations, it is inferred that the
pluton was derived from partial melting of youngsubducted slab material associated with subduction ofthe Kula-Pacific ridge. However, the initial
87Sr/86Sr of
the pluton is 0.70493 ^7, which is higher than MORB
(0.7023–0.7034). Therefore, the magma generation alsoinvolved some high
87Sr/86Sr material. The pluton is poor
in MgO, Ni, and Cr compared with the average adakite,whereas Sr and Na
2OþCaO contents are almost the
same. This indicates that the pluton crystallized from aslightly evolved magma affected by fractional crystal-lization of plagioclase þhornblende. It is concluded that
the pluton was derived from partial melting of a young
subducted slab, contaminated with minor high
87Sr/86Sr
material and affected by fractional crystallization ofplagioclase þhornblende.
Sr and Y data of Cretaceous granitic rocks in south-
western Japan suggested that the granitic rocks were mostlygenerated by partial melting of mafic lower crust underpressures of ,1 GPa. Adakitic plutons derived by slab
melting have also been identified. They are older than thegranitic rocks. Therefore, it is concluded that Cretaceousplutonism in southwestern Japanese arc started with smalladakitic intrusions associated with slab melting, followed by
larger bodies of granite resulting from lower crust melting.
Acknowledgements
This work is based on a portion of the author’s doctoral
dissertation at Yamaguchi University. I am grateful toMasaaki Owada and Yasuhito Osanai for very helpfuldiscussions. I also thank Keiichi Shiraki, Hiroo Kagami,
Nobutaka Tsuchiya, Takuji Hamamoto, Toshihiko Saka-
shima, Teruyoshi Imaoka, Toshisuke Kawasaki, OsamuMaishima, Toru Yamasaki, Tetsuya Kosaka, YukiNakayama, and Shunsuke Jikihara for many commentsand encouragement during this study. Shunso Ishiharapermitted use of unpublished data. I wish to thank YojiMorifuku for assistance with mineral analyses. Detailedcomments by Robert J. Stern and Herve ´Martin significantly
improved this paper. Satish Kumar, Kevin Burke andJennifer Lytwyn carried out important language corrections.
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Earth and Planetary Science Letters 564 (2021) 116893 Contents lists available at ScienceDirect HLAVS Earth and Planetary Science Letters ELSEVIER www.elsevier.com/locate/epsl Evidence for crustal removal, tectonic erosion and flare-ups from the Japanese evolving forearc sediment provenance Frontieercstiufdiilinrycencekuivsitn b Cente forNorth East Asian StudiesTohokuUniversity,9808576,41Kawauchi,Aoba-ku,SendaiMiyagiJapan Department ofEarthcience,raduatechoolofcience,TohokuUniversityJaan SchoolfEarthndanetarycincehIstitutfrsciensearchTe)urtiniversityerthustal éDepartment of GeologicalSciences and GeologicalEngineering,Queen's University,Kingston, Canada ＡＲＴＩＣＬＥＩNＦＯ ＡＲＳＴＲＡＣＴ Article history: Forearc basins preserve the geologic record relating strictly to arc magmatism. The provenance of forearc Received 15 October 2020 sediment can be used to differentiate periods of crustal growth, accretion, and destruction, enhanced Received in revised form 11 March 2021 magmatism, advancing and retreating subduction slabs, delamination, etc. These tectonic systems predict Accepted 15 March 2021 differing degrees of sedimentary reworking of the older forearc units. Additionally, Hf isotopes of zircon Available online 6 April 2021 Editor: R. Hickey-Vargas can be used to evaluate the degree of continental reworking that occurs in the arc system. In this paper, we evaluate the changes in a long-lived subduction system using detrital zircon U-Pb and Hf-isotope Keywords: data from forearc units in northern Honshu, Japan that span in age from the Silurian Period to the detrital zircon present from the forearc provenance of the Japanese subduction system. Our data demonstrate a series geochronology of dominant age peaks (430 ± 20, 360 ± 10, 270 ± 20, 184 ± 12, 112 ± 22, and 7 ± 7 Ma) and a Lu-Hf progressive loss of the older zircon populations. Zircon Hf-isotope data reveal three discrete shifts that crustal growth correspond to differing degrees of isotopic enrichment and correlate with changes in the dominant zircon tectonic erosion age peaks. Additionally, each temporal isotopic shift is associated with isolation of the older sedimentary packages wherein no detrital zircon from the previous stages are observed in subsequent stages. We propose these shifts provide evidence for rapid shifts in arc tectonics including magmatic flare-ups, producing the dominant peaks; protracted tectonic erosion progressively removing older sources of zircon reveals a late Carboniferous event triggering the complete removal of the Precambrian crust; and the Cretaceous melting of the entire Permian arc crust, likely related with the subduction of the mid-oceanic ridge separating the Izanagi and Pacific plates. @ 2021 Elsevier B.V. All rights reserved. 1. Introduction material from the forearc at convergent margins, originally pro- posed in the Japan and Peru Trenches by von Huene and Lallemand The history of the Earth, since the Archean, is carved in the con- (1990). The global budget of continental crust has not been always tinents' crustal records: the earliest rocks differentiated from the neutral and there have been episodes of net growth and shrink- mantle; the origin, evolution of life; the evolution of the magnetic age (e.g. Spencer et al., 2019). Interestingly, during the episodes of field, or the development of superplumes from the core-mantle net growth parts of the crust may have been vigorously destroyed boundary. Unfortunately, this information is disrupted and frag- and vice versa. Documenting the origin and fate of the conti- mented, the crust not only grows but recycles back into the mantle nental crust is a key goal of the Earth sciences to understanding by a variety of processes. The continental crust grows by magma- Earth's chemical evolution and the main tectonic processes operat- tism in arcs (e.g. Spencer et al., 2017), whereas it gets destroyed ing through time. Sialic crustal growth and destruction are crucial primarily in subduction zones through subduction of sediments, to develop accurate plate restorations and paleogeography, which tectonic erosion (a.k.a. subduction erosion, von Huene and Lalle- in turn are the foundation of Earth history, paleoclimatic studies, mand, 1990), and lithospheric delamination sensu lato (e.g. Magni and tectonic research. Without accurate constraints of when and and Kiraly, 2020). Tectonic erosion is the removal of upper-plate where a piece of crust existed such reconstructions are useless for their primary purpose. * Corresponding author. The Panthalassa-Pacific ocean system has been subducting E-mail address: dpastorgalan@gmail.com (D. Pastor-Galan). below today's East Asia-Oceania for at least 500 m.yrs (e.g. https://doi.org/10.1016/j.epsl.2021.116893 0012-821X/ 2021 Elsevier B.V. All rights reserved. D. Pastor-Galan, C.J. Spencer, T. Furukawa et al. Earth and Planetary Science Letters 564(2021）116893 Maruyama and Seno, 1986; Isozaki et al., 2010). The consump- parts of Japan (e.g. Hitachi, Akiyoshi, Ultra-Tamba) show robust co- tion of this superocean formed two of the largest accretionary incidences with North China Craton and at least these areas may orogens of the Phanerozoic where the continental crust grew sig- have originated there (Tagiri et al., 2011; Wakita et al., 2021). Fi- nificantly from ~1000 to 250 Ma: the Central Asian Orogenic Belt nally, some researchers think that the Japanese crust could be a (Jahn, 2004) and ~800 to 250 Ma in the Terra Australis Orogen fragment from NE Gondwana that migrated northwards during the (Cawood, 2005). In contrast, the Japan arc is a narrow strip with opening of the Neotethys (Permian) as an oceanic island arc to fi- a cyclic record fragmented accretionary complexes and blueschist nally collide with both North and South China (Otoh et al., 1990; (s.l.) exhumation during ~500 m. yr. of subduction preserved over Okawa et al., 2013). a Cretaceous granitic crust. Although tectonic erosion has played From late Ordovician, the same subduction regime apparently an important role in the poor preservation of the Japanese base- continued until the present time. Firstly as a continental arc to ment (e.g. Isozaki et al., 2010), this process has been common finally develop a small-scale back-arc basin with a minimal for- in growing orogens (e.g. Stern, 2011). The reasons why the Japan mation of true oceanic crust (Japan Sea) during the Miocene (e.g. arc behaved so differently to its counterparts in the Panthalassa- Maruyama et al., 1997). The only continent-continent collision Pacific subduction zone are poorly known due to the discontinuity, record in Japan is found in the Hida Belt, an allochthonous unit large gaps, and scarcity of its crustal record. was thrusted over pre-Jurassic units during the Triassic collisional Forearc basins preserve the geologic record relating strictly to orogeny that sutured North and South China cratons (e.g. Isozaki, arc magmatism where the bulk of the continental crust is created 1997; Ernst et al., 2007; Fig. 1). The absence of evidence for other and destroyed. Although sediment from the volcanic arc is trans- collisional events supports a Panthalassa-Pacific facing arc from the ported both to the backarc and forearc regions, the forearc is less Ordovician attached to the South China margin (e.g. Isozaki et al. likely to receive detritus from the hinterland as the magmatic arc 2014) rather than a Gondwana derived Tethyan facing arc colliding often forms a continental divide. In this study, we investigate the against China (Okawa et al., 2013). crustal evolution of the northern Honshu arc (NE Japan) from the The geotectonic units of present-day Japan archipelago are di- Silurian to the present day through the provenance of detrital zir- vided into SW and NE Japan by the Tanakura Tectonic line (Fig. 1)) con and their Hf-isotope signature in the South Kitakami forearc SW Japan is characterized by a series of orogen-parallel accre- basin, which contains a close to continuous sedimentary record tionary complexes; the Median Tectonic Line further separates the (cf. Ehiro et al., 2016). Our new results, together with a reappraisal inner and outer zones. The overall structure is a pile of north- of former studies, show a fierce history of periodic flare-ups, tec- rooting subhorizontal nappes, whose older sheets usually occupy tonic erosion, complete removal of the Precambrian crust in the the upper structural positions. Voluminous calc-alkaline granitic late Carboniferous, and total melting of the Permian crust in the i nns s Cretaceous. phic rocks intruded the nappe structure in Cretaceous times, gently folding them to form synform-antiform structures. This magmatic event has been interpreted as a flare-up caused by the subduc- 2. Geological background tion of a mid-ocean ridge (Maruyama and Seno, 1986). In contrast, the nappe distribution in NE Japan is geometrically more com- The Japanese archipelago is a 3000 km long bow-shaped chain plex due to significant structural complications (Fig. 2) together of islands along the eastern margin of Asia preserving at least 500- with a thick cover of Cenozoic volcanic rocks and sediments. Ex- million-year history of subduction processes (e.g. Maruyama et al., posures of pre-Jurassic geotectonic units and Cretaceous batholiths 1997; Isozaki et al., 2010). At present, it is located at the junc- e n e r tions of four distinct plates Eurasian, Pacific, North American, and Abukuma Mountains. Although the subhorizontal nappe-pile struc- Philippine Sea plates (Fig. 1). The Pacific plate is subducting at a ture is completely disrupted by N-S tending high angle faults, all rate of ~8 cm/yr beneath NE Japan, whereas the Philippine Sea lithological components in NE Japan are comparable to those of plate is subducting from SE to NW with ~3-5 cm/yr under SW SW Japan (e.g., Isozaki et al., 2010). Japan along the Nankai trough (trench) and Ryukyu trench off SW n ) Japan. Permian, which included the collision of an oceanic arc; Juras- 2.1. Geological history of Japan sic; and Late Cretaceous to the present (e.g. Isozaki et al., 2010). Several fragments of serpentinized mantle wedge peridotites/ser- The origins of the present-day archipelago are tied to the breakup of the supercontinent Rodinia (e.g. Maruyama et al, 1997; described: ~540 Ma (Oeyama and Miyamori-Hayachine), ~280 Ma q on d u aio e na -rs (Yakuno), Late Jurassic (Mikabu and Horokanai), Late Cretaceous China cratons that rifted apart from Rodinia while Panthalassa (Poroshiri） and Cretaceous-Eocene crystallization ages (Mineoka) ocean (paleo-Pacific) opened (e.g. Isozaki et al., 2010). After an (e.g, Ishiwatari and Tsujimori, 2003). In addition to these rocks, uncertain period of Precambrian evolution, an ocean (perhaps three high-pressure/low-temperature (HP-LT) metamorphic belts the Panthalassa) commenced subduction below what today is the crop-out with ages of 360-300 Ma (Renge), 240-200 Ma (Suo), Japan arc. The oldest subduction-related rocks in both NW and SW and 120-60 Ma (Sanbagawa and Kamuikotan) (Tsujimori and Itaya, Japan arcs are late Cambrian arc-type granitoids and serpentinized 1999). Notably, the Renge and the equivalents HP-LT rocks oc- mantle wedge peridotites with jadeitites (e.g. Isozaki et al., 2015; ss n Tsujimori, 2017). Despite their paucity and dismembered nature, Miyamori-Hayachine units, and 120-60 Ma HP-LT rocks in the they indicate the subduction history of Japan commenced at least o m n mm udi j ~500 Ma. It has been postulated that the proto-Japan arc formed and LP-HT metamorphic rocks in the inner zone (Miyashiro, 1961; part of eastern Cathaysia passive margin (Fig. 1) from late Pro- Fig. 1). terozoic to early Paleozoic until subduction initiated or flipped in A very particular feature of the present-day arc crust in Japan n s s is that it is mostly Cretaceous or younger. Crustal-scale seismic -s u jo s nod n jo o n nn cross-section of sW Japan reveal pre-Cretaceous rocks that occur gest a connection with Cathaysia (e.g. Isozaki et al., 2010; Isozaki, as roof pendant at shallow depths (e.g. Ito et al., 2009). Despite 2019; Wakita et al., 2021). Some authors suggested, however, that the 400 m.yrs of pre-Cretaceous subduction history, pre-Cretaceous 2 D. Pastor-Galan, C.J. Spencer, T. Furukawa et al. EarthandPlanetaryScienceLetters564(2021）116893 A) B) 45 Fore-arc basin Continental fragment Oceanic arc Arc with continental crust HT-LP belt Nemuro younging 40° Accretionary complexes HP metamorphic belts North China JapanSea 40 Hida belt Fig.2 (North China 35 SWJapan Pacific Ocean Yangtze South China craton Suo au Oe:Oeyama UT: UItra Tamba Mz: Maizuru Ry: Sanyo-Ryoke C) Sb: Sanbagawa 30° 30° MTL: Medium Tectonic Line Time TTL: Tanakura Tectonic Line Event EdI.Ca.. ..o... sI .D. .Car. .Pe .J. I.PaN 550 500 450 400 350 300 250 200 150 100 50 Forearc basin South Kitakami - Kurosegawa Massive carbonates Platform AccretedAccreted Accreted Accret. Authocht. Nedamo1 Akiyoshi Mino-Tanba3 Shimanto /Hidaka compl. Allocht. Uitra-Tanba / Nedamo 2 Island arc collision Maizuru (sialic arc) Okhotsk/Nemuro Ophiolite Oeyama/Miyamori/Renge (gabbro+jadeitite) Yakuno Mikabu/Horokanai Poroshiril Setogawa/Mineoka Serpent. melange Oeyama/Renge &Kurosegawa Kamuikotan HP Metamorphism Fuko-pass Renge Suo Kamuikotan/Sanbagawa/Tokoro HT Metamorphism Ryoke/Higo Hidaka I-type Granitoids nunziew ￥ * Sanin S-type Granitoids Sanyo/Ryoke Collisional orogen *Suggested, very lite direct record Hida (North China-South China collision) Fig. 1. A) Location of the main continental blocks and cratons of East Asia (modified after Harada et al., 2021). Early Carboniferous Tsunatori Unit (Uchino, 2021); 2Permian Takinosawa Unit (Uchino, 2021); 3Including N Kitakami, Oshima and S Chichibu; 4Including Kuril and Palau collisions. B) Simplified geological map of Japan based on the Seamless digital geological map of Japan 1: 200,000 (Geological Survey of Japan, 2021). C) Simplified chronology of the main tectonic events recorded in the Japanese active margin. (For interpretation of the colors in the figure(s), the reader is referred to the web version of this article.) plutonic rocks are very scarce. Exposure of ~500-400 Ma granites 2.2. Geology of the South Kitakami Mountains are limited as fragments in the Kurosegawa Belt that is a klippe- like narrow composite unit in the outer zone of SW Japan and The South Kitakami Mountains (SKM hereafter) lie in Tohoku fault-bounded small blocks of the Kitakami Mountains of NE Japan (NE Honshu, Fig. 2) and contain the only relatively thick Paleo- (e.g. Isozaki et al., 2015; Shimojo et al., 2010). The best exposure of zoic and Mesozoic continental margin forearc basin in Japan. In these granitic bodies is the Hikami Granite (Fig. 2 and SF-1) formed the SKM also crop out some of the scarce Cambrian-Silurian arc ~450-440 Ma (e.g. Shimojo et al., 2010; Isozaki et al., 2015). granitoids (Isozaki et al., 2015); a weakly metamorphosed accre- The ~300-270 Ma granites are extremely rare, and ~250-200 Ma tionary complex (the Motai metamorphic rocks, locally blueschist- granitoids only occur within the Hida Belt in central Japan. Sev- facies), which has been compared to the Renge HP-LT metamorphic eral authors, primarily based on detrital zircon U-Pb ages studies, rocks of the SW Japan (Tsujimori and Itaya, 1999); and a supra- speculated that the pre-existing older arc crusts had been signifi- subduction zone ophiolite (Hayachine-Miyamori Complex; Ozawa cantly removed, probably subducted into the mantle, by multiple et al., 2015). The South Kitakami forearc basin (SKFB hereafter) episodes of tectonic erosion (e.g. Suzuki et al., 2010; Aoki et al., represent an independent tectonostratigraphic unit that contains 2012). This secondary disappearance of older crust contributed to a nearly continuous forearc sequence from Silurian to Early Cre- the shortage of information for paleogeographic reconstruction of taceous that lies over the Hikami granite and in tectonic contact Paleozoic and older Japan. with the accretionary and metamorphic units. 3 D. Pastor-Galan, CJ Spencer, T.Furukawa et al. Earth and Planetary Science Letters 564(2021）116893 The SKFB comprises unmetamorphosed shallow-marine Silurian s n to Early Cretaceous strata (Fig. 1). The succession (Ehiro et al., 2016 than 10% were considered unreasonable and these data were ex- and references therein) starts with a basal arkose overlain by Sil- cluded. sHf(t) values were calculated for all data using the 176Lu urian limestone and tuff to Devonian tuff and interlayered mud- decay constant = 1.865 × 10 - 11 yr-1 (Scherer et al., 2001). and sandstone. Those are unconformably overlain by the Early Car- H z/gi (o) ie na nog ie ae na s boniferous interlayered mud- and sandstone with some tuffaceous = 0.282785, 176Lu/177Hf CHUR = 0.0336. Depleted mantle values na 1 = = rocks followed by massive late Carboniferous limestones. Over a minor unconformity, Permian shallow marine clastic strata with volcaniclastics, limestones, and conglomeratic intercalations occur. was used, while for zircon >1,500 Ma, the 207pb/206pb age was The Mesozoic strata (Triassic to lowest Cretaceous) are located in s a sn e a the southern part of the SKM and were deposited in a shallow dards and reduction of results in included in the supplemental File marine or alluvial environment and are mainly composed of clas- SF1. tic rocks in association with rare limestone and tuff. The Mesozoic We have complemented our dataset with 18 Paleozoic and stratigraphy starts over a disconformity with the Paleozoic strata Mesozoic U-Pb detrital zircon samples from Okawa et al. (2013) and contains several minor unconformities (Fig. 2). This succession (16 Samples newly coded OK) and Isozaki et al. (2014) (2 samples ends with a thick volcanic sequence at the Lower Cretaceous (Ehiro newly coded IS; Fig. 2; full site description in SF-1). In addition, et al., 2016) and was intensively intruded during the Aptian-Albian we also show the zircon U-Pb from 8 plutons that intruded the (Fig. 1; e.g. Tsuchiya et al., 2014; Os0zawa et al., 2019). The Aptian- fore-arc basin of SK (Plutons newly coded OsO after Osozawa et Albian Cretaceous granitoids of the SKM show frequently adakitic al., 2019). Five out of these eight plutonic units also include Hf- or shoshonitic composition and ages ranging from 127-113 Ma isotope analyses in zircon: one from Hikami granite and four from (e.g. Osozawa et al., 2019; SF-1). These plutons are slightly older Aptian-Albian plutons (Tono, Hondera, Kesengawa, and Hitokabe; than the equivalent Cretaceous granitoids exposed in SW Japan. Fig. 2; SF-1). We used the software package BAD-ZUPA (SF1-3) to Previous detrital zircon studies in the SKM (Shimojo et al., quantitatively study the dominant zircon populations in the detri- 2010; Okawa et al., 2013; Isozaki et al., 2014) tried to unravel the tal zircon spectra, in addition to kernel density estimations and enigmatic pre-Mesozoic paleogeography of Japan through sediment multi-dimensional scaling (Vermeesch, 2018). BAD-ZUPA is capa- provenance with contrasting results. The studies found a paucity of ble of automatically identifying the statistically significant peaks Precambrian zircon, which corroborates the forearc setting, where and valleys (at a 95% confidence), their most probable age, and the adxa s n ns s a uncertainty of it. prevents paleogeographic correlations. Okawa et al. (2013） sug- Our sample from Hikami Granite (Kita13) displays no inherited gested the affinity of the SKM with Gondwana. In contrast, Isozaki s () 1 m e e et al. (2014, 2015) emphasized similarities of SKM, SW Japan, and is compatible with previously published ages from U-Pb in zircon E Russia with the South China block, supporting the hypothesis of (OSO-08 in SF1, 442.4 ± 9.8 Ma, and 449.2 ± 4.5 Ma, Osozawa a 'Greater South China Craton'. et al., 2019) and other methods (Ehiro et al., 2016 and references therein). Cretaceous granitoids samples from Osozawa et al. (2019: 3. Sampling, methods, and results OSO-01 to OS0-07) show ages between 120 and 110 Ma (see SF1). In general terms, detrital zircon age spectra of samples from Fourteen fore-arc sedimentary clastic samples with ages from the Kitakami forearc present unimodal or bimodal age distribu- Silurian to present-day and one igneous sample (Hikami pluton, tions whose main peak is increasingly young in tandem with late Ordovician-early Silurian) were collected from the SKM of the stratigraphy (Fig. 3). With the single exception of the Orik- NE Honshu (Samples coded Kita; Fig. 2, Supplementary File SF- abetoge formation (Kita17, IS1) the maximum depositional ages, 1). Biostratigraphic constraints demonstrate the age of sedimentary defined by the youngest concordant analysis, are in line with the units spanning from the Silurian Period to the present (Fig. 2 and biostratigraphic constraints (Fig. 3). Silurian and Devonian sam- SF-1 for in detail rock formation). Zircon extraction followed tra- ples (Kita17, 12, 07; OK1-3 and IS1) have a unimodal distribution ditional mineral separation techniques (crushing, milling, sieving, with a late Silurian to early Devonian maxima, similar to the ages Wilfley table, Franz magnetic separation, and heavy liquid sepa- from Hikami granite. Carboniferous samples Kita06 and OK4 are ration). Zircon were mounted in epoxy, imaged with cathodolu- bimodal with the former late Silurian peak and a major Carbonif- minescence, and analyzed for U-Pb and Hf-isotopes during two erous one (~355 Ma) whereas IS2 show a wide unimodal dis- sessions via laser ablation inductively coupled plasma mass spec- tribution with a peak in 356 but including the Silurian-Devonian trometry (LA-ICP-MS) in the John de Laeter Centre (JdLC) at Curtin maxima in the distribution. All Permo-Triassic samples (Kita16, 11, University with a Resonetics RESOlution M-5OA-LR, incorporating 02, 05, and 0K5-10) are unimodal with a Permian (290-260 Ma) a Compex 102 excimer laser. Following a 15-20 s period of back- peak except for OK10 (Late Triassic), which shows a minor Tri- ground analysis, samples were spot ablated for 30 s at a 7 Hz assic peak. Jurassic and Cretaceous samples are unimodal (OK11, ii jo a sse s e sn n Kita04, 15), bimodal (0K12, 13, Kita10, 01), or multimodal (0K14 at the sample surface. The sample cell was flushed by ultrahigh and 16). The Permian peak is represented in all the samples and - ( 8) 7 1 (1 ) Ad an early Jurassic one (~190-180 Ma) is common to all but OK11 ties were measured using an Agilent 770Os quadrupole ICP-MS and and Kita04. The youngest Cretaceous sample (OK16) shows a Creta- a Nu Instruments Plasma I MC-ICP-MS, with high purity Ar as the ceous peak (~130 Ma). The Cenozoic samples present a prominent plasma gas (flow rate 0.98 L min-1). On the quadrupole, most ele- Cretaceous peak (~105 Ma) with ages similar to the SKM Creta- ments were monitored for 0.01 s each with the exception of 88Sr Ie) qdgoz 9dzoz '9d9oz '9doz (s t0'0) 1d11 (s v0'0)e16e1 (s 200) Pb 0.03 s) 232Th (0.0125 s), and 238u (0.0125 s). Approximately not relevant, 12 samples contained 0 Precambrian zircon, and the half of the split was sent to a Nu Plasma II MC-ICP-MS for Lu- others just a few, which never cluster. These Precambrian zircon + g i ss -aueinus painseau aiaM JHost Pue 'JH6zi JHgLi 'JHzi 'nT + qX spectra and a composite spectrum with all samples (Figs. 3 and s a s o on n sod us ( calculated ages from two U-Pb systems lie within uncertainty of from the oldest representative peak. This is especially noticeable in 4 D. Pastor-Galan, C.J. Spencer, T. Furukawa et al. Earth andPlanetaryScienceLetters564(2021）116893 Clastics Limestones Accretionary complex Tuffaceous clastics M Unconformity 10 km Nedamo Belt (Carbiniferous) Grain size Permian onglomerate North Kitakami Belt (Jurassic) KitaTsu1&2 ·Kita 08 i5. OSO-1-7 Cretaceous plutonism 2 Kita 01& OK16 39.5° 39.5° ta17 Kita 10&OK15 Tono Pluton : 8K14 OS0-1 Kita04&OK12 Hitokabe Granite ·OK11 OSO-2 ·OK10 Goyo-san granite 2 S-01 OSO-3 Kita 05&OK9 Kita07 ita12 Kita 02 & OK8 &Kita08 OK7 · OK06 Kita 11 Hikami Granite Kita06 39° &0K4 OSO-8 275 Hirota Pluton ·Kita 16 OSO-7 HonderaPluton Kita 06 OSO-4 Kesengawa OK5 Pluton OS0-6 OrikabeGranite OK8 OSO-5 3 OK5 OK11 3 .Kita 06 & OK4 OK12 OK.13 ·IS2 South Kitakami Forearc basin ●Kita07&OK3 Pre-Carboniferous .35 Kita10&OK15 Carboniferous OK7 Permian 38.5° Triassic Jurassic Ao Cretaceous Granitoids and volcanics Kita02 Kita01 Silurian granitoids ·Kita 17&IS1 Kita05 Cretaceous granitoids Cretaceous volcanics Hikami Granite Kita04 Metamorphic rocks ·OK1&OK2 Paleozoic ●Kita 13& OSO8 KitaTsu 1&2 OK14 Hayachine-Miyamori ophiolites Hayachine-MiyamoriophiolitesM Cambrian-Ordovician OK16 Fig. 2. South Kitakami synthetic geological map and stratigraphy showing the sampling locations based on the Seamless digital geological map of Japan 1: 200,000 (Geological Survey of Japan, 2021). Samples coded Kita are newly analyzed, OK after Okawa et al. (2013), IS after Isozaki et al. (2014) and samples coded OS0 after Osozawa et al. (2019). D. Pastor-Galan, CJ Spencer, T.Furukawa et al. EarthandPlanetaryScienceLetters564(2021）116893 KDEwith95%Cl Biostratigraphicage A) ● Youngest concordant zircon Best peak age Zircon forming event n = Phanerozoic zircons PreC. Kitatsu2 n = 40 LOO Kitatsu1 n = 35 Kita08 n = 41 8000 Kita01 n =68 0 100 200 300 400 500 1000 2000 3000Ma OK15n=19 2 Pre-Permian samples n=695 B) Kita10 n=60 100 9 gra Kita04n 2 58 69 n=1094 300 OK9n=69 200 Kita05n=70 OK8n ：113 Detrital zircon Kita02n=50 sourceremoval OK7n=91 Cenozoic samples n=115 ：54 OK6n 304050 Kita11 n = 40 Kita16n=35 20 Detrital zircon 10 source removal OK4n=65 age [Ma] Kita06n=76 0 100 200 300 400 500 /S2n=55 Fig. 4. A) Composite spectra of all detrital samples. B) Composite spectra of all OK3 n = pre-Permian samples. C) Composite spectra of the Permian-Mesozoic samples. D) Spectra of the three Cenozoic samples studied. Kita07n=64 Kita12n=79 A composite age spectrum of all detrital samples defines the n od n Kita17 n =35 ber of pre-Cambrian zircon is residual (n = 190, 11%), none of the /S1n=54 peaks is statistically relevant. If we consider only the total Pre- OK2 n=54 cambrian zircon, 3 peaks are significant: ~0.6 Ga, 1 Ga, and 1.9 OK1n=58 Ga. We identified 6 synthetic dominant zircon populations at 430 Age 100 ± 20, 360 ± 10, 270 ± 20,184 ± 12,112 ± 22, and 7 ± 7 Ma0 200 300 400 500 Ma. We have plotted all individual samples in a multi-dimensional Fig. 3. Zircon spectra in the Kita (new dataset), OK (Okawa et al., 2013) and IS scaling (MDS) map (Vermeesch, 2013) against and the synthetic (Isozaki et al., 2014) samples from the South Kitakami forearc basin. zircon population ages (Fig. 5; Spencer and Kirkland, 2016). MDS transforms a matrix of pairwise similarities (the D value from the Permian-Mesozoic samples, which have no or negligible amounts showing all detrital zircon populations considered. On a MDS dia- of pre-Permian zircon; and Cenozoic samples, whose spectra dis- gram distances represent the degree of similarity, the smaller the play no zircon older than ~120 Ma. distance between two samples the more similar they are (Ver- 6 D. Pastor-Galan, C.J. Spencer, T. Furukawa et al. Earth and Planetary Science Letters 564(2021）116893 2+L O112±22 (Figs. 3 and 6). A synthetic detrital zircon spectrum considering all Kitais02 new and literature samples from SKM has 6 statistically significant populations at 430 ± 20, 360 ± 10, 270 ± 20,184 ± 12, 112 ± 22, and 7 ± 7 Ma (Figs. 3, 4 and 5). All samples but Orikabetoge KitaTs01 184±12 the new data, 400.2 ± 7.9 Ma, for phengite separates (8.139 ± 0.16 wt Kita08 line with the biostratigraphic age (Figs. 2 and 3). The youngest con- of Late Paleozoic detrital zircons from sedimentary (and metasedimentary) rocks from the Cathaysia Block of South China Craton [SC] (Hu et al., 2012),the Yeongnam Massif [YN] (Cheong 360±10 oKita11 published biostratigraphic age (~30 m.yrs). Its zircon distribution is, nonetheless, very similar to the other Silurian units. Both sam- 2 ·OK5 at ~440 Ma. Early Paleozoic calc-alkaline granitoids with zircon U-Pb OK16. Kita16 mentions the lithological similarities between Silurian and Devo- .OK15 0 Kita04. nian clastic rocks in that area and geologic relationships are often OK13 270±200 OK8 obscured by minimal rock exposure (Ehiro et al., 2016). Therefore, OK12 IS2· ng Kita058 OK7 OK140 we cannot rule out that samples were collected in a Devonian unit. -0.2 ·OK4 Kita18 Kita06° OK6 Kita120 OK11 4.1. Provenance of the South Kitakami forearc 430±20 OK9 -0.4 The sedimentary system of the Japanese forearc in SKM expe- oIS1 rienced limited sedimentary reworking of older forearc material and little sediment support from cratonic and orogenic areas lo- T -0.5 0.0 0.5 cated to the west (present-day coordinates). The major peaks in each sample (Fig. 3; Okawa et al., 2013; Isozaki et al., 2014) devi- Fig. 5. Multidimensional Scaling map of all detrital zircon samples depicting the ate very little from the dominant synthetic zircon age populations similarities and evolution of the three identified groups based on the total loss of the previous distributions (pre-Permian, Permian-Mesozoic and Cenozoic). phism in proto-Japan is poorly understood due to the paucity of Early 7 Ma). Additionally, samples fall into three categories defined by meesch, 2013). The plot is dimensionless and values range be- activities and granulite-facies metamorphic rocks would have formed tween 0, and 1 on each axis. A distance of 0 between two samples Permian-Mesozoic, and Cenozoic. The zircon support to the basin is initiation (Tsujimori, 2017; Tsujimori & Harlow, 2017). Early Paleozoic controlled therefore by several detrital zircon forming events (co- butions). The MDS map displays three clusters: the Pre-Permian boundaries (late Carboniferous-early Permian and late Cretaceous) samples; the Permo-Mesozoic samples and the Cenozoic samples. The i76Hf/177Hf isotopic signature in zircon represents a proxy when the older arc and forearc stopped supporting zircon to the to estimate when the rock that crystallized such zircon was ex- basin. The paucity of Precambrian zircon (190/1991, ~10%) indicates tracted from the mantle and to diagnose crustal reworking through time, where successive samples define a Hf evolution array (e.g. that the volcanic arc has acted as a long-term barrier impending geotherm. The occurrence of jadeitite associated with serpentinite de- transport of zircon from any craton. The combined Precambrian (Kita13 and Os0-08) have initial εHf values of -10 to +1 (Fig. 6A). age spectra show three main populations (Fig. 6B) at ~600, ~1000, The εHf values for the SKM Cretaceous granitoids (OsO-01 to 07) and 1900 Ma, and two minor peaks at ~1500 and ~2800 Ma, vary from 5 to 13, although other Tohoku areas (see Osozawa et which are consistent with a minor inflow of sediment to the fore- al., 2019) exhibit a wider range (from -20 to 15), being from 0 to arc basin from South China Craton where the proto-Japan arc was 15 the most concentrated area (Fig. 6A). Precambrian zircon show likely located until the Triassic (Isozaki, 2019). So far, the 1900 Ma very variable eHf signatures, ranging from -20 to positive values peak has not occurred in NE Gondwana, and the ~10o0 Ma one close to the depleted mantle curve (Fig. 6B). No zircon older than is largely absent in North China (e.g. Zhao et al., 2017). The disap- grown rim of bright luminescence (L11 of Fig. 6). Three spot analyses on pearance of Precambrian zircon during the Permian-Triassic times lution through the stratigraphy shows an eHf increase from ~430 and reappearance in the SKM record from the Jurassic might be Ma, with initial eHf values very similar to those of Hikami gran- indicative of the Permo-Triassic collisional events in east-central n 0 1 1 ' China (e.g. Isozaki, 1997; Ernst et al., 2007) and its sedimentary ~360 Ma, eHf trend increases with a less pronounced slope but losing all the less juvenile sources to ~270 Ma, where values get from after such collision. close to the depleted mantle curve. From there eHf decreases until Despite the SKFB preserves an almost continuous stratigraphy ~112 Ma following a typical crustal residence trend. At around 112 from Silurian to Cretaceous, the prospective pre-Cretaceous arc- Ma detrital zircon register, as in the igneous rocks, values ranging related sources of zircon to the basin are generally lacking. Apart from quite positive to about O, mimicking the Cretaceous grani- Does the Kitomyo Schist provide a clue for the first generation of toids trend. The Hf-isotope array displays a similar effect in the plutonic rocks older than Cretaceous are extremely scarce in SKM sub recent population of zircon (Fig. 6B). in particular and Japan in general (Fig. 1). In the SKM, the Silurian Hikami Granite (and equivalent plutons not cropping out) could 4. Discussion represent the main source for the 430 ± 20 population. The only preserved putative sources for the 360 ± 10 population are mi- The new samples (Kita) and reappraised (OK, IS and OsO sam- nor tuffs (Fig. 2; SF1). No sources for 270 ± 20 and 184 ± 12 Ma ples from Okawa et al., 2013; Isozaki et al., 2014 and Osozawa et populations have been identified so far in SK, being the nearest in al., 2019, respectively) provide a crucial source of information to age and location the Carboniferous-Permian Wariyama granite in understand the Phanerozoic crustal history of Japan. The detrital the Abukuma Mountains, to the south (~300 Ma., Tsutsumi et al., zircon spectra through the stratigraphic column of SKFB revealed 2010; Tsuchiya et al., 2014). The Permian zircon population is the Zone metamorphism. On the other hand, no Early Paleozoic blueschist- most prominent not only in SK, but also in other in other Permo- 391 Ma. If we consider discordant data between the two rim ages and Mesozoic forearc basins and accretionary complexes in SW Japan tratigraphic ages; and with minor to no Precambrian contribution Ito, & Tamura, 2020; Tsujimori, 2010; Tsujimori & Itaya, 1999; Tsujimori 7 D.Pastor-Galan,C.J. Spencer,T.Furukawa et al. Earth and Planetary Science Letters 564(2021) 116893 AI ●Hikami OSO ■Kita13 Depleted Mantle Cretaceous OSO Other Cretaceous H3 nitial 10 0.012 77Hf -10 -20 Age (Ma) 100 200 300 400 500550 1550 2550 3550 ○Kita01Kita08 20 B Kita02Kita10 Kita04 Kita11 Kita05 Kita12 VKita06 ●Kita16 Kita07Kita17 KitaTs01 KitaTs02 10 -10 Age(Ma) 0 100 200 300 400 500550 155025503550 Fig.6. A) eHf in zircon respect to their U-Pb zircon age for the Hikami Granite (Silurian) and the Cretaceous granitoids of South Kitakami and surroundings (from Osozawa et al., 2019) B) eHf in the detrital zircon vs. their ages (Kita samples). neous rocks are almost absent in the Japanese record excepting isotope array (Fig. 6B). Zircon of the ~1.9 Ga population and older the Hida Belt. In contrast, Miocene to recent andesites and rhy- show variable negative values of eHf. The younger populations olitic flows and, especially, Cretaceous large batholiths abound in (~1 Ga and ~600 Ma) have very mixed values, from very posi- all the Japanese archipelago (Figs. 1 and 2). tive to negative, suggesting a mixing between significant amounts of newly extracted from the mantle material and other crustal 4.2. Hf isotopes: Japan sinks, Japan melts sources. Our Precambrian results are compatible with a provenance from Cathaysia and/or Yangtze blocks (South China craton, Fig. 1), The Precambrian population of zircon in Kita samples is too with similar detrital zircon populations and εHf array (Cawood et scarce (n = 49) to interpret important trends in the crustal res- al., 2018; Wan et al., 2019). In contrast, North China's 1.9 Ga popu- idence of the source areas. Nonetheless, we think it is a useful lation has quite positive εHf values (Xia et al., 2008). Further data preliminary constraint into the contested proto-Japan paleogeog- is necessary to confirm such affinity, but with the present dataset, raphy (South China vs. North China, e.g. Isozaki, 2019; Wakita et we are inclined towards a proto-Japan arc being part of a Greater al., 2021). We distinguished two groups in the Kita samples Hf- South China continent (Isozaki et al., 2014, 2015). 8 D.Pastor-Galan, CJ. Spencer, T.Furukawa et al. Earth and Planetary Science Letters 564(2021）116893 The evolution of Hf-isotope in the Phanerozoic zircon through but one. Older populations get progressively less important, to the stratigraphy (Fig. 6) shows a Hf-isotope array that starts with finally disappear forever (Fig. 3). In addition to the progressive similar values to the Hikami granite (Kita 13 and OsO 08 from loss of previous sources, both the individual and the composite Osozawa et al., 2019). The age similarity and εHf signature sug- gest that Hikami and/or very similar non-preserved Silurian plu- pre-Permian zircon age peaks (Figs. 3 and 4). Likewise, Cenozoic tons fed the forearc basin during the Silurian-Carboniferous times. The eHf/Ma trend displays progressively more juvenile (positive) and 4). This leaves three groups (Pre-Permian, Permian-Mesozoic, values as the zircon population gets younger. This indicates that and Cenozoic, Fig. 4) that are further justified by the use of MDS the sources of Devonian and Early Carboniferous zircon mixed de- by comparing the SKFB samples with synthetic age populations pleted mantle material with the previous SKM crust. From ~360 (Fig. 5). In the MDS map, the Pre-Permian samples cluster near Ma, the trajectory keeps on increasing until ~270 Ma with very the early Silurian population with a trend as samples go upwards positive values in the proximity of the depleted mantle curve. Re- in the column towards the early Carboniferous. Following a simi- markably, the trend from 360 to 270 Ma loses most of the less lar pattern, the Permian-Cretaceous samples cluster near the ~270 juvenile contribution, indicating little crustal mixing of the sources ± 20 Ma synthetic population with a trend towards the Jurassic of Carboniferous and Permian zircon. We interpret that the steep maxima (184 ± 12 Ma). Finally, the Cenozoic samples cluster near εHf/Ma trajectory from 450 to 270 Ma implies a progressive loss the ~112 ± 22 Ma but not far from the 7 ± 7 Ma population. of the original proto-Japan crust. At the beginning of such a pro- sr ssnd s e s two main events where previous sources disappear could not be d m p n the same. 1 SKFB contains an almost continuous forearc stratigraphy from dicating an almost complete loss of the previous crust during the late Silurian to early Cretaceous (Fig. 2), and most of other ge- Late Paleozoic. This process ended with a complete crustal replace- ological evidence in the archipelago points to an uninterrupted ment in the Permian. subduction below Japan during, at least, 400 m.yr (e.g. Maruyama From ~270 Ma, εHf trend decreases until ~112 Ma following a et al., 1997; Isozaki et al., 2010). If arc activity had been contin- typical crustal residence trend. We interpret the trend as a period uous and approximately at a constant rate, we would have found in which the Permian new crust matured and where, despite the a progressively younger maximum depositional age pointing to a Jurassic minor flare-up, new mantle input was minor. At around progressively younger arc source feeding the basin. However, we 112 Ma detrital zircon register, analog to the igneous rocks, val- found an age consistency of the dominant populations through ues ranging from quite positive to about O, the same as in the the stratigraphy, suggesting that significant zircon forming events sub recent population of zircon (Fig. 6B). It indicates a lot of mix- P ing between the Permian crust and new inputs from the depleted magmatism. We propose that the main populations represent mag- mantle. Considering that the majority of the present-day Japan arc crust is a Cretaceous large batholith, we support that the major- eval to main HP-LT metamorphic events (Fig. 1): (1) Silurian - Fuko ity of the Japanese crust melted during a punctual and diachronic Pass; (2) early Carboniferous - Renge; and (3) the most obvious, event (in Tohoku ~112 Ma, progressively younger to the South and Cretaceous - Sanbagawa, where Cretaceous batholiths paired with west, Osozawa et al., 2019). Some authors attributed the magmatic coeval HP-LT. The cyclic HP-LT metamorphic together with coeval event to the subduction of the Izanagi plate's ridge below Japan LP-HT metamorphism and anatexis (Miyashiro's ‘paired metamor- (e.g. Maruyama and Seno, 1986; Maruyama et al., 1997), although phic belt' concept (1961) are common in long-lived Pacific type the kinematics, timing, and orientation of the ridge subduction are orogeny (e.g. western USA: Snoke and Barnes, 2006) and poten- disputed (see Wu and Wu, 2019 for discussion and references). tially related to flare-ups. The only tectonic events close to coeval The gap in the SKFB zircon record from the Cretaceous popula- to the Permian maxima is the collision of Maizuru arc and the sub- tion to the recent population hinders the crustal evolution during recent population fits in time with the opening of the Japan Sea the Late Cretaceous and Paleogene. Other fragments of the fore- (Fig. 1). We could not find any event that can explain the Jurassic arc basin like the Izumi Group in SW Japan) may shed light on zircon maxima. the crustal evolution since samples generally show unimodal de- The progressive depletion of older populations upwards in the trital zircon peak at ~80 Ma, however, Hf-isotope signatures have stratigraphy indicates the loss of the igneous source and litle re- not been studied so far (Aoki et al., 2012). The sub-recent pop- working of previous strata. Low rates of basin reworking suggest ulation shows a range of juvenile εHf (from ~0 to ~12). This protracted subsidence keeping the basin away from erosion. The indicates, again, the mixing of new depleted mantle sources and fading of the older zircon sources may be related to its burial be- former crust. In this case, we think that the retreat of the arc after low newer arc material; high erosion rates enough to completely the Miocene opening of the Japan Sea (Van Horne et al., 2017 and erode the source rocks; or tectonic erosion removing the oldest references therein) explains best the Miocene to recent Hf-isotope section of the arc from below. Complete burial would require con- trend. tinuous arc production. Our zircon record evidences intermittent magmatic flare-ups instead. Large amounts of magmatism during a 4.3. Flaring up and dragging Japan down flare-up could be blamed for the burial of older sources. In such a case, it would be expected that these older sources became more The Phanerozoic zircon pool of individual sample spectra shows present after several million years of erosion of the flare-up, but dominant populations that are roughly coeval and little Precam- we found the opposite. The complete erosion of previous arcs man- brian zircon sources (Fig. 3). The dominant populations, confirmed ifests similar setbacks: if the arc had been repeatedly dismantled by the composite Phanerozoic detrital zircon spectra of SKFB (1801 due to high erosion rates, we would have found Precambrian zir- out of 1991 zircon), are late Silurian, early Carboniferous, Permian, con coming from the cratons each time the arc did not represent early Jurassic, Aptian-Albian, and sub-recent in age. The youngest a sedimentary barrier. Previous studies (e.g. Suzuki et al., 2010; Isozaki et al., 2010) suggested that the Japan arc has undergone as we go upwards in the stratigraphy (Fig. 3), or not to a sig- frequent periods of tectonic erosion since the Silurian, remov- nificant level. For example, OK6 (Permian strata) has a younger ing significant parts of its geological record. Tectonic erosion (von peak than all Triassic samples whose youngest peak is Permian Huene and Lallemand, 1990) can explain both the disappearance 9 D. Pastor-Galan, C.J. Spencer, T. Furukawa et al. Earth and Planetary Science Letters 564(2021）116893 of older arc sources, the mixing of Hf-isotope signatures becom- A late Carboniferous loss of the Japanese crust, which we interpret ing progressively more juvenile, and continuous subsidence of the as a delamination process, and a Cretaceous melting of the entire forearc. arc, probably related to the subduction of the mid-oceanic ridge Nonetheless, the main two events when the older sources com- separating the Izanagi and Pacific plates. pletely disappeared (Late Carboniferous - Early Permian and Cre- taceous) are hardly explained by tectonic erosion. During the late CRediT authorship contribution statement Carboniferous event, the Hf-isotope array shows the arc crust was completely renewed (Fig. 6B). Delamination of the lithospheric Daniel Pastor-Galan: Conceptualization, Validation, Formal Anal- mantle (e.g. Magni and Kiraly, 2020) can explain the partial re- ysis, Resources, Investigation, Visualization, Data Curation, Writing, moval of the lower crust and the generation of mantle-derived Funding Acquisition. magmatism, whereas the consequent uplift produces rapid and in- Christopher J. Spencer: Conceptualization, Methodology, Formal tense denudation. We think this is the most plausible mechanism Analysis, Investigation, Visualization, Data Curation, Writing-review explaining the quick removal of the Precambrian crust in NE Japan and editing, Funding Acquisition. and the subsequent Permian flare-up (Figs. 3 and 6B). Candidates Tan Furukawa: Software, Formal Analysis, Visualization, Re- triggering delamination in the overriding plate are a rapid arc re- sources. treat, for example, due to slab roll-back, or collision of an arc or Tatsuki Tsujimori: Validation, Writing-review and editing, Fund- oceanic plateau that stagnated below Japan, resulting in an over- ing Acquisition. thickened crust and forming a convective drip at the base of the thickened lithosphere. So far, we have not been able to find com- Declaration of competing interest pelling evidence of late Carboniferous widespread extension and no The authors declare that they have no known competing finan- gin (e.g. Shen et al., 2018). In contrast, the only collisional candi- date (Maizuru; Fig. 1) is too young and it is uncertain whether the cial interests or personal relationships that could have appeared to magnitude of the collision was enough to precipitate a delamina- influence the work reported in this paper. tion event (Fig. 1, Isozaki et al., 2010). The Hf-isotope signature of the Cretaceous zircon sources loss event suggests, in contrast to Acknowledgements the late Carboniferous, a general mixing between the older and the juvenile crust. The crustal reworking time is coeval to the We would like to thank Xiaofang He, Brad McDonald, and Noreen Evans for lab assistance. Hiroyuki Okawa and Soichi Os- composition (Tsuchiya et al., 2014; Osozawa et al., 2019) that rep- ozawa kindly provided their datasets to be included in this paper. resents the bulk of today Japanese crust, and the final subduction Research in the GeoHistory laser ablation Facility, John de Laeter of the Izanagi plate (c.f. Wu and Wu, 2019). We found appeal- Centre, Curtin University of Technology is supported by AuScope ing a cause-effect relationship between the final subduction of the (auscope.org.au) and the Australian Government via the National Izanagi ridge, the formation of slab windows, and the development Collaborative Research Infrastructure Strategy (NCRIS). The NPII of a flare-up that adds new material from the depleted mantle and multi-collector was obtained via funding from the Australian Re- was capable of melting the majority of the Permian crust of Japan. search Council LIEF program (LE150100013). The constructive com- We also think that the Miocene opening of the Japan Sea explains so e on well the character of the sub-recent zircon population. The detrital N Ssi A Pap sm ded s aded a Pad zircon of SKFB tell us the hidden history of the missing record of Grant (JP15H05212, JP16F16329, JP18H01299) and an “Ensemble Japan, a violent tale of flare-ups, crustal melting, and lithospheric Grant for Young Researchers" (Tohoku University). DPG also thanks foundering. a lifetime of company provided by Christopher John Cornell. You broke your rusty cage, now the cold Earth is your bed. 5. Conclusions Appendix A. Supplementary material The composite Phanerozoic detrital zircon spectra of the SKFB of Japan define dominant zircon age populations at ~430, ~360, Supplementary material related to this article can be found ~270, ~180, ~112, and ~7 Ma. We found very few Precambrian online at https://doi.org/10.1016/j.epsl.2021.116893. These data in- zircon, which indicate the arc acted as a barrier for craton sup- clude the Google map of the most important areas described in port during all Phanerozoic. In addition to these dominant popu- this article. lations, we recognize a progressive disappearance of older sources and three distinct time intervals with very few to no zircon: ~320 a1e sde asaui eW 0z 0n 09~ pue eW 0t1 0 09I~ :eW 00e 01 References matism. The 176Hf/177Hf isotopic signature of the zircon spectra Aoki, K, Isozaki, Y., Yamamoto, S., Maki, K., Yokoyama, T, Hirata, T., 2012. Tectonic erosion in a Pacific-type orogen: detrital zircon response to Cretaceous tectonics shows an εHf increase from ~430 Ma to ~270 Ma when the in Japan. Geology 40, 1087-1090. https://doi.org/10.1130/G33414.1. Japanese crust became completely juvenile. From there eHf de- Bouvier, A., Vervoort, J.D., Patchett, PJ., 2008. 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LITHOS 404-405(2021)106440 Contents lists available at ScienceDirect OHL LITHOS ELSEVIER journal homepage: wv .com/locate/lithos Research Article Petrogenesis of amphibole-rich ultramafic rocks in the Hida metamorphic complex, Japan: Its role in arc crust differentiation Osamu Ishizuka é, Akihiro Tamura a a School of Geosciences and Civil Engineering, College of Science and Engineering Kanazawa Universty, Kanazawa, Ishikawa 920-1192, Japan olcIeel 0061, Japan d School of Environmental Science and Management, Universty of the Philippines Los Banos, Laguna, Philippines ARTICLEINFO ABSTRACT Keywords: Amphibole controls major and trace element chemistry of arc magma via fractionation, therefore its formation Amphibole process is the key to understand arc crustal evolution. Amphibole-rich cumulates, occasionally found associated Arc magma with felsic plutons or metamorphic rocks, provide complementary insights into the arc magma differentiation. In Hornblende peridotite this study, we conducted a petrological examination of amphibole-rich ultramafic rocks from the Hida meta- Crustal differentiation morphic complex in Japan to better understand its role in arc crust evolution. The studied samples are char- Lower crust acterized by large poikilitic amphibole and orthopyroxene enclosing olivine grains with corroded margins. Amphibole showed high Cr203 and TiO2 contents (up to 0.9 and 1.2 wt%) and light rare earth element enrichment relative to heavy rare earth elements. Microscopic and geochemical features of amphibole and olivine suggest that the amphibole formed at the expense of early crystalized olivine (± clinopyroxene) from hydrous juvenile magma at a deep crustal level, which is consistent with the estimated crystallization pressure of 0.7-1.0 GPa based on an empirical Al-in-hornblende geobarometer. Considering the arc affinity of parental magma and the amphibole-fractionation trend in bulk chemistry of the Mesozoic plutonic rocks from the Hida Belt, it is implied the formation of amphibole-rich plutonic rocks contributed to arc crustal evolution as amphibole-rich residue complementary to differentiated magma. 1. Introduction ascent of low-T silicic magma produced by fluid-fluxed melting of mafic underplates and to contribute to the production of low-K and calk- Arc magmatism has been regarded as a fundamental process in alkaline granitoid suites (Collins et al., 2020). generating continental crust with a differentiated and andesite compo- Amphibole has also been regarded as an important fractional phase sition, which is unique to Earth within our solar system (Rudnick, 1995; for arc magma differentiation in the last several decades (Davidson Tatsumi et al., 2015; Taylor and McLennan, 1996). It has been tradi- et al., 2007; Hidalgo and Rooney, 2010; Kratzmann et al., 2010). The tionally thought that amphibole crystallization causes an efficient in- rare earth element (REE) compositions of arc volcanic suites show global crease in SiO2 content in differentiated magma relative to olivine + trends of increasing La/Yb and slightly decreasing Gd/Yb ratios clinopyroxene fractionation (Barclay and Carmichael, 2004; Cawthon accompanied by increasing SiO2 content, indicating that the critical et al., 1973; Cawthorn and O'Hara, 1976; Holloway and Burnham, phase in arc magma differentiation is amphibole rather than garnet 1972). The role of amphibole in arc magma differentiation has recently (Davidson et al., 2007). This amphibole-fractionation signal can be drawn attention again (Alonso-Perez et al., 2009; Jagoutz et al., 2011; observed even when amphibole phenocryst is uncommon in arc lavas, Mintener and Ulmer, 2018; Nandedkar et al., 2016, 2014). Amphibole which is called “cryptic" amphibole fractionation (Davidson et al., 2007; is also expected to be an important residual/cumulative phase during an Kratzmann et al., 2010; Larocque and Canil, 2010). * Corresponding author. E-mail address: k.itano@staff.kanazawa-u.ac.jp (K. Itano). https://doi.org/10.1016/j.lithos.2021.106440 Received 28 January 2021; Received in revised form 26 August 2021; Accepted 28 August 2021 Available online 2 September 2021 0024-4937/@ 2021 Elsevier B.V. All rights reserved. K. Itano et al. LITHOS 404-405 (2021) 106440 Amphibole-rich plutonic rocks are essential in better understanding 1.1. Geological outline the amphibole formation at a lower crustal level and provide petrolog- ical evidence that it represents possible counterparts of the differenti- The Hida metamorphic complex, one of the oldest basement rocks in ated arc magmas. Amphibole-rich cumulates are not much in volume but Japan, is exposed on the north side of southwest Japan (Fig. la). The are ubiquitous as a part of arc crustal section in convergent margins and Hida Belt is regarded to have been the eastern part of the North China orogens: the Ryoke complex, Japan (Kutsukake, 1978); the Adamello Craton (Isozaki et al., 2010; Takahashi et al., 2018), whereas the tec- batholith, Italy (Ulmer et al., 1983); the Bonanza arc, Canada (Larocque tonic development of the Hida Marginal Belt is controversial (Kunugiza and Canil, 2010); the Chelan complex in the Washington Cascades, and Maruyama, 2011; Tsujmori, 2002; Tsujimori et al., 2006). The western U.s. (Dessimoz et al., 2012); the Glenelg River Complex in the development of the metamorphic complex in the Hida Belt was related Delamerian Orogen, Australia (Kemp, 2004); Husky Ridge locality in the to the Dabie-Sulu orogeny caused by the Mesozoic collision between the Ross Orogen, Antarctica (Tiepolo and Tribuzio, 2008). The number of North and South China Cratons (Arakawa et al., 2000; Sohma and detailed petrological and geochemical studies of the amphibole-rich Kunugiza, 1993), but the geological correlation between the Hida Belt ultramafic rocks is still limited despite their importance due to their sr s r d sporadic occurrences. Gyeonggi massif (Horie et al., 2010; Ishiwatari and Tsujimori, 2003) and In this study, we examine the formation conditions of the amphibole- Ogcheon Belt (Takahashi et al., 2018, 2010; Takehara and Horie, 2019). rich ultramafic rocks (hornblende peridotite) from the Hida meta- The plutons associated with this metamorphic complex are mainly morphic complex in Japan focusing on the mineral chemistry of granitic in composition (Fig. 1b), but the occurrence of syn-plutonic amphibole, olivine, and orthopyroxene. This study demonstrates an mafic intrusions have also been reported (Ando and Ogasawara, 1998; example of effective amphibole crystallization from primitive arc Arakawa et al., 2000). magma at a lower crustal level and highlights the significance of Metamorphic rocks in the Hida terrain are divided into two parts: amphibole-rich ultramafic rock formation for Mesozoic crustal evolu- low pressure/temperature (P/T) Hida metamorphic complex and me- tion of the East Asia continent. dium P/T Unazuki schists (Fig. 1b). The Hida metamorphic complex is (a) (b) 40° North China Hida Belt block Hida Granites Hida Belt (Early Mesozoic) Hida Metamorphic rocks Unazuki Metamorphic rocks Hida Marginal Belt Metamorphic & /MedianTectonicLine Sedimentary Rocks South China block 30° 120° 130° 36° 50km138° Mafic gneiss (Hida metamorphic rocks) Felsic gneiss Calcareous gneiss Quaternary volcanic rocks (andesite to dacite) Cretaceous sedimentary rocks (Tedori group) Cretaceous volcanic rocks (Nohi rhyolite) ++ Triassic granitic rocks (Hida older granite) Sampling point 2km Fig. 1. (a) Map of the Japanese arc showing the position of the Hida metamorphic complex. (b) Simplified geological map of the Mesozoic Hida Belt and Paleozoic Hida Marginal belt. The inset shows the location of the sampling locality in the western part of the Hida metamorphic complex. (c) Sampling locality in the Hokuriku area. Solid lines represent faults. These maps are modified after Horie et al. (2010) and Takahashi et al. (2010) and reference therein. GM: Gyeonggi massif, OB: Ogcheon Belt. K.Itano et al. LITHOS404-405(2021)106440 composed of paragneiss, orthogneiss, calcareous gneiss, and amphibo- Early Mesozoic granitic rocks in the Hida terrain are classified into: lite. These gneiss layers intercalate each other and dominant lithology (i) older (Triassic) granites exhibiting gneissose structure and (i) varies from one locality to another (Arakawa et al., 2000; Asami and younger (Jurassic) granites showing no deformed structure (Kano, Adachi, 1976; Kano, 1991). The protoliths of the gneisses were thought 1990). Zircon geochronology has confirmed that the age gap between to have been various, although the lack of cherts in addition to bimodal the deformed granites of 250-230 Ma and the undeformed granites of 200-180 Ma (Horie et al., 2018; Takahashi et al., 2018, 2010; Takehara shelf sediments were part of the protolith (Sohma and Kunugiza, 1993). and Horie, 2019; Zhao et al., 2013). The metamorphic history of the Hida metamorphic complex was not The outcrop of an ultramafic lens was found in the western part of - a - a the Hida metamorphic complex (36°15'11.6"N, 136°44'20.9"E, Fig. 1b, isochron ages (Arakawa, 1984; Arakawa et al., 2000; Asano et al., 1990); c). The ultramafic body is surrounded by the older granite and mafic gneiss concomitantly formed in Triassic, although its contact boundary trometry (SIMS) revealed that the regional metamorphism forming the was not determined due to dense vegetation and slope reinforcements around the outcrop exposure. The exposure is a small body that is less et al., 2000; Takahashi et al., 2018; Takehara and Horie, 2019). The than 10 m wide. No difference in texture and lithology was observed Archean and Proterozoic zircon inheritances in the gneisses corrobo- between the accessible outcrop and adjacent boulders. Three represen- rated the relation between the Hida belt and North China Craton (Sano tative samples (HU0526 1-b, 3-a, and 4-b) were used for detailed et al., 2000; Takahashi et al., 2018), and the inherited zircon ages of petrography and geochemical analyses in this study. igneous cores also constrained the maximum depositional age of pro- tolith to >252 Ma (Horie et al., 2018; Takehara and Horie, 2019). 1 cm Amph 50 μm 1mm 100 μm B 1 mm Amph mm Fig. 2. Thin section photographs and back scattered-electron (BSE) images of hornblende peridotites from the Hida metamorphic complex. (a) Cross polarized light image. Large poikilitic amphibole grains showing the same extinction angle are bordered by lines. (b) Poikilitic orthopyroxene. Amphibole (purplish grain in upper left part) enclosing olivine chadacryst is further enveloped by poikilitic orthopyroxene. (c) Poikilitic amphibole showing euhedral faces in contact with the orthopyroxene (gray grain in central part) enclosing olivine. (d) BSE image of spinel showing chemical zoning. (e) Plane-polarized light image. Phlogopite occurs as an interclumulus phase. (f) BSE image of magnetite (bright parts) in serpentine veinlets. (g) Al elemental map of amphibole acquired by EPMA. The greenish part corresponds to pargasitic composition, whereas the bluish part corresponds to tremolitic composition. Red circles and numbers represent analytical spots and id for of the references to colour in this figure legend, the reader is referred to the web version of this article.) K. Itano et al. LITHOS 404-405 (2021) 106440 1.2. Samples and petrography and BCR-2 were used for quality control (Supplementary Data Table S1). Major element compositions of rock-forming minerals were deter- The studied samples are characterized by a poikilitic cumulative mined with an electron probe micro-analyzer (JEOL JXA-8800 Superp- texture with olivine chadacryst enclosed by large amphibole or ortho- robe) at Kanazawa University. This analysis was conducted using an pyroxene oikocrysts (>1 cm, Fig. 2a) and classified as hornblende accelerating voltage of 20 kV, a beam current of 20 nA, and a spot size of peridotite (Table 1). Amphibole-rich ultramafic rocks with such peculiar 3 μm for the core and rim compositions of olivine, amphibole, ortho- texture used to be referred to as cortlandites (Williams, 1886). The pyroxene, and spinel. The measurement data were calibrated and cor- hornblende peridotite and/or hornblendite showing similar poikilitic rected using natural and synthetic minerals and the ZAF correction scheme. In-house mineral standards (olivine, chromian spinel, diopside, (Larocque and Canil, 2010; Tanaka et al., 1982; Ulmer et al., 1983). and K-feldspar) were repeatedly measured to check the data accuracy. Minor and accessory phases include phlogopite, spinel, magnetite, The trace element compositions of orthopyroxene and amphibole apatite, Fe—Ti oxide, and pyrite. It is noteworthy that clinopyroxene and plagioclase are absent in the studied samples, although these min- single-spot ablation of 50 or 100 μm in diameter at a 6 Hz repetition rate erals are commonly present in amphibole-rich plutonic rocks as minor with an energy density of 8 J/cm? and NIST-SRM 612 was used as an phases. external standard. The Si contents measured with EPMA were adopted Isolated olivine grains of several mm in size are rounded or have as internal standards for all trace element analyses. We carefully resorbed textures, suggesting a reaction with the enclosing phases selected analytical spots to match those to the analytical spots for EPMA (Fig. 2b, c). Olivine is basically enclosed by amphibole first (Fig. 2b) and data and confirmed data matching based on monitoring of Si and Ca signal intensities in LA-ICPMS data (e.g., Fig. 2e). NIST-SRM 614 glass is the largest oikocryst (up to 2-3 cm) and euhedral facets along was also analyzed for quality control during measurement, which orthopyroxene contacts were observed (Fig. 2c). A textural relationship showed identical values with reference values within 10% relative wherein amphibole and orthopyroxene include each other obviously standard deviation for all the analytical elements. Details of this indicates the same crystallization stage. The exsolution lamellae of analytical procedure are described in detail by Morishita et al. (2005). ilmenite are occasionally found and such regions were avoided for chemical analyses. Cr-rich spinel is presumed to be a primary phase that 2. Results commonly occurs as inclusions mainly in olivine and amphibole, and rarely in orthopyroxene. The size of spinel ranges from 50 to 300 μm and 2.1. Whole-rock chemistry shows a clear chemical zoning (Fig. 2d). Phlogopite occurs as an inter- clumulus phase and its shape is controlled by the original outlines of Whole-rock geochemical data are summarized in Table 2. The olivine grains (Fig. 2e). Serpentine and magnetite are post-igneous analyzed samples are ultramafic (~43 wt% Sio2) and homogenous in phases. Olivine grains are partly or considerably serpentinized along major and trace element compositions. The average trace element the cracks, and serpentine occurs as magnetite-rich or magnetite-free composition of the studied samples is shown in the pyrolite normalized veinlets (Fig. 2f). Most magnetites are present in serpentine veinlets or diagram (McDonough and Sun, 1995) (Fig. 3). The studied samples are occurring as clots (Fig. 2f). Metamorphic alteration such as chlorite is characterized by an enrichment of light-REE relative to heavy-REE (La/ generally very minor. We mainly focus on the early-magmatic mineral Yb ～ 4.0 and Dy/Yb ~ 1.7) and by the depletion of Nb, Ta, and Ti assemblage (olivine, orthopyroxene, and amphibole) in this study to relative to other elements of similar incompatibility (Fig. 3). The trace reveal magmatic processes responsible for the formation of the horn- element pattern of the studied samples is similar to that of average arc blende peridotite. basalts rather than mid-ocean ridge basalt (MORB) (Kelemen et al., 2003), although the abundances are lower than those of the arc basalt (Fig. 3). 1.3. Analytical methods 2.2. Mineral composition Whole-rock analyses for major and trace elements were carried out at Kanazawa University. The major element analysis was conducted by X- Representative data are summarized in Table 3 and all data are ray fluorescence (XRF, ZSX primus II) according to the method described in Kusano et al. (2014). The trace element analyses were carried out by (Mg + Fe2+)) of olivine ranges from 83.4 to 84.7 (83.8 on average) and laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP- Ni content slightly varies from 0.12 to 0.18 wt% (Table 2, Table S2), MS) using a 193 nm ArF Excimer laser system (MicroLas GeoLas Q-plus) which deviate from the mantle olivine array (Arai, 1987; Takahashi coupled to a quadrupole ICP-MS (Agilent 7500S). A 100 μm laser beam et al., 1987). Orthopyroxene has higher Mg# (85-86) compared to with an energy fluence of 8 J/cm? at a repetition rate of 5 Hz was used, olivine (Fig. 4a). The Al2O3 and Cr2O3 contents of orthopyroxenes are up and BCR-2G was used as an external standard. The fused glasses of to ~3.6 wt% and ~ 0.35 wt%, respectively (Table S2). Fig. 5a shows the samples were prepared by a direct fusion method using an iridium-strip comparison with Mg# of olivine and orthopyroxene and the average heater for the trace element analysis and the details are described in value of this study (olivine: ~84 and orthopyroxene: ~86) lies on the Tamura et al. (2014). This method is simple and quick compared to a equilibrium line defined by literature data (Gasparik, 1984; Matsui and conventional solution nebulizer-ICP-MS method, although it is difficult Nishizawa, 1974; Ozawa, 1986; Zeck et al., 1982), which indicates the to accurately determine the concentration of volatile elements (e.g., Li, Mg—-Fe re-equilibrium between olivine and orthopyroxene at subsolidus B, Cl and Pb) due to loss during heating. The fusion glasses of BHVO-2 condition. The amphibole structural formula was calculated by assuming a Table 1 stoichiometry (Li et al., 2020) and fall under the calcic group with Petrographic characteristics of hornblende peridotite. compositions ranging from pargasite to tremolite following the Samples Mineral proportions nomenclature scheme of Hawthorne et al. (2012). The Mg# of amphi- bole calculated using total iron (100 Mg/(Mg + FeTotal) shows higher HU0526 1-b Amph (44), 01 + Spt (44), Opx (5), Phl (3), Opq (2), Spl (1) HU0526 3-a Amph (45), 0l + Spt (41), Opx (10), Phl (1), Opq (1), Spl (1) Mg# of 84-91 and its large variation relative to other minerals (Fig. 4a). HU0526 4-b Amph (46), 0l + Spt (43), Opx (6), Phl (2), Opq (2), Spl (1) The variation in amphibole Mg# would reflect the tschermak substitu- Amph, amphibole; Ol, olivine; Spt, serpentine; Opx, orthopyroxene, Opq, opa- tion causing an increase in Mg2+ and si4+ associated with a decrease in que minerals; Phl, phlogopite; Spl, spinel A13+ (Table S2). Considering the linear correlation between Si cation K. Itano et al. LITHOS 404-405 (2021) 106440 Table 2 and Mg# for calcic amphiboles (Fig. 5b), the Mg# of magmatic Whole-rock geochemical data of hornblende peridotite. amphibole that started to crystalized from primitive magma would be HKC1-1 HKC1-2 ~86 at most (Krawczynski et al., 2012). The chemical composition of Sample: HKC1-3 o dn s o pe sr uu aq pazno ose si asrd wt% 43.53 43.32 43.11 0.9 wt% and 1.2 wt%, respectively. SiO2 TiO2 0.58 0.47 0.44 Spinel shows intra-grain and inter-grain variations in chemical Al2O3 5.78 6.54 6.54 compositions (Table 2). The rim domains yielded lower Cr# (Cr/ Fe2O3 13.48 13.16 13.66 (Cr + Al)) and higher Mg# than the core domains (Cr# core - Cr#rim: 0.07 MnO 0.20 0.18 0.20 in average, Mg#core - Mg#rim: -0.1 in average; Fig. 4b). This chemical MgO 29.85 29.60 29.47 CaO 4.10 4.00 4.17 zoning is explained by the transfer of Al and Mg from the surrounding Na2O 0.78 0.74 0.78 silicate minerals to the spinel during subsolidus exchange (Bai et al., K20 0.41 0.55 0.47 2018; Cameron, 1975; Henderson and Suddaby, 1971), which is P2O5 0.06 0.06 0.05 consistent with the loss of Al in magnesio-hornblende and tremolite Total 98.77 98.62 98.88 compared to magmatic pargasite. No correlation between spinel ppm 3.7 9.0 5.2 composition and the type of host mineral including spinel was observed. Sc 20.5 17.5 17.6 The chondrite-normalized REE and primitive mantle-normalized Ti 3553 2723 2586 trace element patterns of amphibole and orthopyroxene are shown in V 147 120 122 Fig. 6, and all trace element data are reported in Supplementary Data Cr 1074 1030 1334 Co 110 124 Table S3. The pargasite composition is characterized by the slight 120 Ni 741 824 797 enrichment in light-REE to heavy-REE (Fig. 6a) and the significant de- Rb 4.5 7.4 9.1 pletions in Nb, Ta, Zr, Hf, and Ti (Fig. 6b). The trace element pattern of Sr 123.1 128.6 118.9 amphibole is identical with the whole-rock pattern because amphibole is Y 11.2 10.9 10.1 Zr 29.0 39.3 42.7 the host phase of trace elements in the mineral assemblage. On the other Nb 1.15 1.27 1.23 hand, the co-existing orthopyroxenes exhibit positive Zr, Hf, and Ti Cs 0.38 0.72 1.72 anomalies (Fig. 6b). Ba 74.3 93.9 80.3 La 4.23 4.75 4.41 Ce 10.4 12.2 11.9 Id 1.28 1.53 1.50 PN 5.92 6.91 6.75 Sm 1.62 1.74 1.65 Eu 0.505 0.530 0.519 Gd 1.88 1.85 1.73 (a) Tb 0.304 0.291 0.271 Dy 2.00 1.90 1.79 06 Ho 0.411 0.402 0.365 Er 1.24 1.20 1.10 #8 Tm 0.173 0.169 0.151 1.13 1.14 1.06 M Yb Lu 0.180 0.189 0.189 86 H 0.904 1.10 1.12 0.082 0.116 0.116 Th 0.798 0.811 0.799 84 U 0.144 0.151 0.177 Olivine Orthopyroxene Amphibole 100 Average composition 0.5 Spinel ●Hornblende MORB ■ Continental arc basalt ←core peridotite ■ Oceanic arc basalt rim e 0.4 0.3 # 10 centrat 0.2- 0.1 (b) 0.0 0.4 0.5 0.6 0.7 Mg# Fig. 3. Pyrolite-normalized trace element patterns of average whole-rock samples. The average composition of MORB (Sun and McDonough, 1989), Fig. 4. Major element compositions of minerals. (a) Comparison of Mg/ oceanic and continental arc basalt (Kelemen et al., 2003) are shown for com- (Mg + Fetotal) for olivine, orthopyroxene, and amphibole. Data points are parison. Normalization values are from McDonough and Sun (1995). plotted in violin-plot style and the width in the x-axis direction graphically shows the approximate frequency of data points in each group. (b) Relationship of Mg/(Mg + Fe2+) atomic ratio (Mg#) and Cr/(Cr + Al) atomic ratio (Cr#). K. Itano et al. LITHOS 404-405 (2021) 106440 0.92 (b) Fig. 5. (a) Mg—Fe partitioning rela- The review of the evolution of ideas on the tectonic evolution of ene. Literature data are from Gasparik, 0.90 a common regional origin of those blocks as Jiamusi was maybe Ozawa, 1986; Zeck et al., 1982. (b) #6W #6W (1） The structure of SW Japan is made of a pile of sub-horizontal (2) The mechanisms advocated for the tectonic building within Olivine Amph 0.88 blue circles represent pargasite, mag- For the Permian-Triassic Akiyoshi orogeny (Fig. 8), the scenario 60 0.86 Pargasite would approach theprimary This study Magnesiohornblende magmatic compositions.(For inter- 0.84 pretation of the references to colour in Literature OTremolite this figure legend, the reader is referred to the web version of this 60 80 100 6.2 6.4 6.6 6.8 7.0 7.2 7.4 article.) OpxMg# Si cation Table 3 Representative major element compositions of major phases in hornblende peridotite. Mineral Olivine Amphibole Orthopyroxene Spinel Phlogopite Sample ID Grain-1 core Grain-1 rim Amph12-core Amph178 Grain-2 core Grain-2 rim Grain-2 core Grain-3 core Grain-1 core Grain-1 rim 40.4 40.1 44.5 43.6 56.2 55.9 38.8 38.4 UDL 800'0 1.25 2.20 0.076 0.025 0.13 0.04 0.32 85'0 UDL UDL 14.1 13.3 1.99 2.71 9 36.8 41.1 17.2 16.2 CrO UDL UDL 0.463 0.579 0.186 0.119 23.71 0.022 0.142 FeOtotal 15.6 15.8 5.00 5.60 9.63 9.72 24.04 22.64 6.27 6.51 MnO 0.19 0.21 0.09 0.34 0 0.23 0.242 0.226 0.18 0.17 MgO 45.6 45.3 17.8 17.4 32.0 31.6 9.74 10.59 23.68 24.29 CaO 0.006 0.006 11.6 11.4 0.337 0.178 1 0.06 0.10 NaO 2.48 2.56 - 0.086 0.063 KzO 0.39 0.36 6.92 6.42 NiO 0.12 0.17 0.06 0.05 200 00 8000 0.027 600 0.073 Total 101.9 101.6 97.7 97.4 100.7 100.6 98.35 98.30 93.57 92.99 Si 0.997 0.994 6.37 6.24 0.988 0.987 6.26 6.21 Ti 000'0 0.14 0.26 100°0 000°0 000°0 0.001 0.04 20°0 A1 2.27 2.24 1.303 0.551 1.422 1.63 1.54 Cr 20°0 20°0 100°0 100°0 0.650 0.00 10'0 Fe2 0.321 0.327 Fe3+ 0.45 0.35 0.142 0.143 0.039 0.022 0.85 0.88 0.23 0.32 0.565 0.535 Mn 0.004 0.004 0.01 0°0 00°0 000 900°0 0.006 0.03 0.02 Mg 1.679 1.676 3.64 3.70 0.838 0.831 0.436 0.464 5.70 5.85 Ca 0.000 0.000 1.71 1.76 900°0 Na 000 0.01 0.02 0.66 0.71 0.01 10'0 K 0.07 0.07 0.71 0.66 Ni 0.002 3.005 00'0 0.01 0.01 000°0 000°0 0.001 0.001 0.01 10'0 Total 3.003 15.64 15.75 2.000 1.999 3.002 3.002 15.2 15.3 Mg# 83.9 83.7 86.3 84.6 85.6 85.3 0.44 0.46 0.87 0.87 Cr# 0.28 UDL, Under detection limit; Mg#, 100 Mg/(Mg + Fe); Cr#, 100 Cr/(Cr + Al). 3. Discussion reported signature of primary amphibole crystallized from primitive arc magma (Conrad and Kay, 1984; Dessimoz et al., 2012; Larocque and 3.1.Parental melt composition Canil, 2010). The amphibole crystallization from primitive magma is also sup- The poikilitic texture and the euhedral facets of amphibole (Fig. 2c) ported by trace element chemistry. The partitioning coefficients be- suggest a crystallization from melt, and the igneous phases (olivine, tween amphibole and melt for high field strength elements (HFSE) orthopyroxene, and pargasite) show high Mg# of ~8684 (Figs. 4a and highly depend on melt composition and temperature (Tiepolo et al., 5b). Even if we take the subsolidus FeMg re-equilibrium into consid- 2007). Hf and Zr are incompatible in amphibole for hotter mafic melt eration (Fig.5a), the high Mg# is a remarkable feature of the studied (> ~ 900 °C); in contrast, Hf and Zr are compatible with colder felsic samples. Amphibole with high Mg# of >80 have been reported in high- melt at >890830 °C (Nandedkar et al., 2016). The Hf and Zr in the pressure crystallization experiments of water-saturated basaltic or high- amphibole grains show a negative correlation with Ti that is always Mg andesitic melts (Krawczynski et al., 2012; Pichavant and Macdonald, compatible with amphibole (Fig. 7). Given that amphibole was an early 2007) and natural amphibole crystallizes from high-Mg basalt (Ishimaru crystallization phase and the effect of accessory minerals was not sig- et al., 2009), whereas more differentiated melt such as andesite and nificant during the early crystallization stage, the correlations in Fig. 7 rhyolite crystallizes amphiboles with lower Mg# of <80 in a wide range suggest that Ti behaves as a compatible element and that Zr and Hf of pressure and temperature conditions (Grove et al., 2003; Sisson and behave as incompatible elements during amphibole crystallization. We Grove, 1993). Furthermore, the high Cr and Ni contents up to ~6000 carried out trace element modeling of amphibole fractional crystalliza- and 700 ppm, respectively, in amphibole are consistent with the tion, where we used the partitioning coefficients of Nandedkar et al. K. Itano et al. LITHOS 404-405 (2021) 106440 (a) · Pargasite ·Magnesiohornblende o Tremolite · Opx respectively calculated assuming Rayleigh fractionation. The slopes of obtained evolution trends are roughly consistent with those of obser- vations, although abundances are dependent on initial compositions. Therefore, it is very likely that the amphibole crystallization was related 101 to primitive magma rather than low-temperature differentiated melt. C icentration 3.2.Amphibole crystallization mechanism 10° The poikilitic textures in the studied samples indicate the early for- mation of olivine chadacrysts (Fig. 2b, c). Moreover, a mutual including suo relationship of amphibole and orthopyroxene oikocrysts suggest their 8 almost simultaneous crystallizations. The trace element chemistry of the 10-1 melts in equilibrium with amphibole and orthopyroxene also suggests a common source (Fig. 8). The REE composition of melt in equilibrium La Ce Pr NdPimSmEu Gd Tb Dy Ho Er Tm Yb Lu with amphibole was calculated using the average pargasite composi- tions and amphibole-melt partitioning coefficients for REE (Shimizu (b)101 et al., 2017) at temperatures ranging from 830 to 1010 °C because it is PM difficult to determine the crystallization temperature for the studied mineral assemblage. Likewise, the equilibrium-melt compositions for uo orthopyroxenes were calculated using the partitioning coefficients be- rat 10° tween orthopyroxene and hydrous basalt (McDade et al., 2003). Patterns u from both of the melts in equilibrium are characterized by light-REE ce enrichment relative to heavy-REE. The REE concentrations of equilib- 10 rium melts are highly dependent on temperature, however, the changes Co in relative ratios of DREEs (i.e, the shape of pattern) are negligible on a chondrite-normalized logarithmic scale. Hence, we conclude that the Rb Th Nb La Sr Nd Hf Eu Ti Dy Ho Tm Lu amphibole and orthopyroxene formed from the same parental melt based on trace element chemistry and textural relationship. Olivine, included amphibole and orthopyroxene, shows rounded or Fig. 6. Trace element compositions of amphibole and orthopyroxene. resorbed textures (Fig. 2b, c). This textural disequilibrium is commonly Chondrite-normalized values are shown for (a) rare earth elements and Primi- observed as a result of the incongruent reaction between hydrous tive mantle-normalized values are shown for (b) analyzed trace elements. The basaltic melt and olivine and clinopyroxene (Beard et al., 2004; Caw- REE concentrations of amphiboles vary depending on major element compo- thorn and O'Hara, 1976; Foden and Green, 1992; Nandedkar et al., sitions. Normalization values are from Sun and McDonough (1989). Data with 2014; Sisson and Grove, 1993; Smith, 2014). Therefore, amphibole values below the limit of detection are excluded from the graph (#202 and crystallized at the expense of early-crystallized/existed olivine. #311 in Supplementary Data Table S3). One question that needs to be asked is what is the requirement for the formation of amphibole-rich plutonic rocks (>45 vol% amphibole). Smith (2014) highlighted the important role of clinopyroxene pre- 65 O [wdd] H O cursors for efficient amphibole formation. They demonstrated the nat- Zr [ppm] ural evidence of clinopyroxene-replacement reaction forming -2.0 60 % crystallization amphibole based on textural relations (e.g., amphibole bleb in F=60% Z 55- O O 1.7 △Opx 70% 生 B 50 O O Pargasite 80% 102 1.5 45- 90% /CI 40- 3000 4000 5000 6000 1010°℃ M1011 [udd] ! Fig. 7. Trace element chemistry of magmatic amphiboles. Solid lines represent various degrees of fractional crystallization. Partitioning coefficients between amphibole and calc-alkali magma of Nandedkar et al. (2016) are used. A melt in equilibrium with the amphibole (Amph-18 in Table S2) is assumed as the initial 10° 830°C △ composition and then Hf, Zr, and Ti contents in fractionated melts and equi- librium amphibole are calculated using the Rayleigh equation. Hafnium and La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu zirconium contents show negative correlations with titanium content and these observed trends are consistent with the calculated evolution curves. Fig. 8. Rare earth element compositions of melts in equilibrium with magmatic amphibole and orthopyroxene. Amph/LD values for 830-1010 °C used in the calculation and warm colour refers to hotter conditions (For interpretation of (2016) and assumed a melt in equilibrium with the amphibole (Amph- the references to colour in this figure legend, the reader is referred to the web 18 in Table S2) as the initial composition. The compositions of frac- version of this article). Partitioning coefficient source: Shimizu et al. (2017) and tionated melts and equilibrium amphibole compositions were McDade et al. (2003). K. Itano et al. LITHOS 404-405 (2021) 106440 clinopyroxene) and trace element chemistry; however, no clinopyroxene the eastern margin of the Asian continent (~30 km) (Chisenga and Yan, was observed even as an inclusion in all our samples. Amphibole shows the LREE-rich REE patterns (Fig. 6a) and this is inconsistent with the (Isozaki et al., 2010; Liu and Chen, 2017; Sano et al., 2000; Takahashi HREE-rich REE pattern expected for amphibole replacing clinopyroxene et al., 2018). When we assume the formation of the studied hornblende (Smith, 2014). Indeed, hornblende-rich cumulates having no clinopyr- peridotite took place around the Moho depth, it is plausible to explore the contribution of the amphibole-rich cumulative rocks to the Mesozoic Canil (2010) and the high-Cr hornblende oikocryst is interpreted as the crustal evolution. involvement of a primitive magma for amphibole crystallization. Hence, The trace element composition of melts in equilibrium with the average pargasite, calculated using amphibole-melt partitioning co- primitive magma at the consumption of olivine (± clinopyroxene) is the efficients of Nandedkar et al. (2016), shows geochemical similarity to key to understanding the petrogenesis of our samples, although we average continental arc basalt (Fig. 9). The enrichments of light-REE, cannot rule out the process of clinopyroxene consumption. It is also Th, and U in the equilibrium-melt are interpreted as a crustal affinity implied that the replacement of clinopyroxene is not necessarily and suggest the contribution of crustal components to the source mantle required for the formation of amphibole-rich plutonic rock. via slab fluid/melt (Plank and Langmuir, 1998). The negative Nb-—Ta anomalies in addition to this enrichment are generally observed in primitive arc basalt (Kelemen et al., 2003). High Nb/Ta ratios of par- 3.3.Amphibole crystallization condition gasite (~20) also implies that the efficient amphibole crystallization affected the Nb/Ta fractionation in the differentiated magma. Amphi- The early crystallization of amphibole at high temperature occurs bole fractionation during cooling of arc hydrous magma controls Nb/Ta under high-pressure conditions of >0.7 GPa for both hydrous basalt and fractionation in arc magma at crustal levels (Li et al., 2017), whereas primitive andesite systems (Foden and Green, 1992; Grove et al., 2003; rutile saturation dominantly controls Nb/Ta fractionation in mature Pichavant and Macdonald, 2007). In such hydrous systems, plagioclase continental arcs with >50 km crustal thickness (Tang et al., 2019). regardless of melt compositions (Barclay and Carmichael, 2004; Grove Considering the formation depth of the hornblende peridotite at 20-30 km, the amphibole fractionation would have played an important et al., 2003; Pichavant and Macdonald, 2007). On the other hand, garnet role in Nb—Ta systematics. These trace element signatures are in can be stable at high pressure of >1.2 GPa in the hydrous systems agreement with the arc magma signature calculated using previously (Alonso-Perez et al., 2009; Mintener et al., 2001). The absence of reported amphibole data in hornblende-rich rocks (Tiepolo et al., 2012, plagioclase and garnet in the mineral assemblages of studied samples 2011; Tiepolo and Tribuzio, 2008) (Fig. 9). Although there are slight suggests that the crystallization of hydrous melt formed amphibole-rich variations in mineral assemblages (±clinopyroxene, ±orthopyroxene, ultramafic rock at moderate pressure conditions (~0.5-1.2 GPa). ±plagioclase), textures (poikilitic or idiomorphic), between hornblende- Amphibole geobarometer verified the moderate formation pressure rich rocks, their geochemical affinities suggest the involvement of that was suggested by the phase relationship. In the high thermody- similar arc magma to form hornblende-rich rocks. namic variance system (i.e., mineral assemblages consisting of a few The genetic relation between the amphibole-rich residue and differentiated rocks was implied by a comparison with geochemical data calibrated barometer on the basis of AlV atoms in amphibole (Krawc- of contemporaneous plutonic rocks in the Hida Belt (Fig. 10). Since the zynski et al., 2012; Larocque and Canil, 2010). We used the following hornblende peridotite is enclosed within Jurassic gneiss and granite, we equation of Krawczynski et al. (2012): P (GPa) = 1.675*A1V-0.048 ± 0.15 (1) 1000 The calculated octahedrally coordinated Al atoms in the amphibole O Melt in equiribrium with pargasite based on normalization to 23 oxygens are represented by A1Vi in this Average continental arc basalt s e o n s assd Tiepolo & Tribuzo (2008) Mg# of <86 was obtained using this equation. Although the geo- 100 Tiepolo et al. (2011) barometer includes an uncertainty reaching ±0.15 GPa due to the MORB ▲Tiepolo et al. (2012) regression of experimental data (Krawczynski et al., 2012) and consid- ering the effect of temperature (Anderson and Smith, 1995), the results show that the studied hornblende peridotite has formed at a deep crustal 之 10 level (~20-30 km). Melt/ 3.4. The role of amphibole-rich cumulates for the crustal evolution of the Hida Belt The above discussions led to the following conclusion on the for- mation of hornblende peridotite. The hydrous primitive magma crys- 0.1 tallized amphibole and orthopyroxene while consuming earlier ThUNbTaLaCesrNd ZrHfsmTiYb crystallized olivine (± clinopyroxene) at lower crustal level (~20-30 km depth). The intrusion of the hydrous primitive magma formed the Fig. 9. N-MORB normalized trace element compositions of melt in equilibrium hornblende peridotite with poikilitic texture. This study supports the with pargasite in the hornblende peridotite and average continental arc basalt (Kelemen et al., 2003). Partitioning coefficients and Normalization values are view that voluminous amphibole crystallization to form amphibole-rich plutonic rocks does not necessarily require highly differentiated magma from Nandedkar et al. (2016) and Sun and McDonough (1989), respectively. Equilibrium melt is chemically similar to the average continental arc basalt, (Larocque and Canil, 2010; Smith, 2014), and this formation process although Th and U abundances in the equilibrium melt are higher than those of could have played an important role in the crustal evolution of the Hida the arc basalt. Equilibrium melt compositions using previous reported brown Belt because of the large impact of the hornblende crystallization from a amphibole data from amphibole-rich rocks are present for comparison: Tiepolo primitive magma on magma compositions. et al. (2011); (2012), Tiepolo and Tribuzio (2008). (For interpretation of the The estimated formation depth of the studied hornblende peridotite references to colour in this figure legend, the reader is referred to the web (20-30 km) roughly corresponds to the base of continental arc crust of version of this article.) K. Itano et al. LITHOS 404-405 (2021) 106440 3.0 2.5 口 1.5 Amphibole 1.0 40 50 60 70 SiO, [wt%] Plutonic rocks (Gabbro ~ Granite) Hornblende peridotite/Hornblendite Arc volcanic rocks ●Hida Belt ■Hida Belt Other areas Japanese Arc Other areas Fig. 10. Dy/Yb vs. SiO2. The whole-rock data for hornblende peridotite and hornblendite are from this study in addition to Larocque and Canil (2010) and Smith n s nq u n ond e d a q m a p (to relation as well as the trend of arc volcanic rocks, which is explained as amphibole fractionation trend by Davidson et al. (2007). The data for cotemporaneous plutonic rocks are from Arakawa and Shinmura (1995), Arakawa et al. (2000), and Ishihara (2005), and those for arc volcanic rocks are from Davidson (1987), George et al. (2004), Handley et al. (2007), and Haraguchi et al. (2018). chose Triassic plutonic rocks in the Hida Belt for comparison. The hornblende peridotite in the Hida Belt. The poikilitic amphibole and studied hornblende peridotite and other plutonic rocks show Dy/Yb orthopyroxene crystallized from the same primitive hydrous magma trend in Fig. 10, which is generally observed for volcanic rocks of the while the early crystalized cumulative phases such as olivine (± clino- Japanese and other arcs. These general trends for arc volcanic rocks pyroxene) were consumed. The crystallization pressure of amphibole were interpreted as the geochemical fingerprint of amphibole fraction- was estimated to be ~0.7-1.0 GPa based on the amphibole geo- barometer, corresponding to a deep crustal level of matured arc. The major and trace element compositions of amphibole also suggested that trend of Dy/Yb for plutonic rocks of the Hida Belt suggests that the amphibole was crystallized from hydrous primitive magma rather amphibole fractionation was the dominant process for magma differ- than already differentiated magma. The Dy/Yb versus SiO2 relationship entiation and the contribution of garnet-bearing lower crustal rocks was for bulk rock compositions of the plutonic rocks suggest the amphibole minor during the flare up producing Jurassic plutonic rocks of the Hida fractionation dominantly contributed to the arc magma differentiation Belt. Hence, the studied hornblende peridotite can be regarded as one of and the crustal evolution of the Hida Belt. the petrological evidences of amphibole-rich residue complementary to Supplementary data to this article can be found online at https://doi. more evolved contemporaneous magma in the Hida Belt org/10.1016/j.lithos.2021.106440. The sporadic occurrences of similar hornblende-rich rocks in the East Asian Continent are interesting because the formation of hornblende- Funding rich rocks might have globally contributed to the crustal evolution of the East Asian Continent. Mesozoic hornblende-rich plutonic rocks were S P e a S s S also reported from Gyeonggi massif in the Korean Peninsula (Kim et al., lows (Grant No. JP19J00913) and research grants from Hakusan 2011; Yi et al., 2016), which is a possible continuation of the Hida Belt Tedorigawa Geopark (issuance year 2018 and 2019). during the Mesozoic metamorphism and plutonism (Arakawa et al., 2000; Ishiwatari and Tsujimori, 2003). Thus, further research should be done to comprehensively investigate the hornblende-rich plutonic rocks Declaration of Competing Interest in the East Asian Continent. The authors declare no conflict of interest. 4.Conclusions Acknowledgments We conducted petrological and geochemical examinations of the This paper was siginificantly imporved by Madhusoodhan Satish- K.Itano et al. LITHOS 404-405 (2021) 106440 Kumar and an anonymous reviwer. This work was supported by a Grant- Grove, T.L., Elkins-Tanton, L.T., Parman, S.W., Chatterjee, N., Muntener, O., Gaetani, G. A., 2003. Fractional crystallization and mantle-melting controls on calc-alkaline differentiation trends. Contrib. Mineral. Petrol. 145, 515-533. https:/doi.org/ [Grant No. JP19J00913]. This work is also supported by research grants 10.1007/s00410-003-0448-z. from Hakusan Tedorigawa Geopark, Japan [issuance year 2018 and Handley, H.K., Macpherson, C.G., Davidson, J.P., Berlo, K., Lowry, D., 2007. 2019]. Constraining fluid and sediment contributions to subduction-related magmatism in indonesia: ijen volcanic complex. J. Petrol. 48, 1155-1183. https://doi.org/ 10.1093/petrology/egm013. 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Journal of Asian Earth Sciences 219 (2021) 104888 Contents lists available at ScienceDirect Journal of Asian Earth Sciences ELSEVIER journal homepage: www.elsevier.com/locate/jseaes U-Pb ages and sandstone provenance of the Permian volcano-sedimentary sequence of the Hida Gaien belt, Southwest Japan: Implications for Permian sedimentation and tectonics in Northeast Asia Keisuke Suzuki ? *, Toshiyuki Kurihara aGraduateSchoolof Science andTechnology,Nigata University,Nigata 9502181,Japan PDepartment of Geology,Faculty of Science,Nigata University,Nigata 9502181,Japan ARTICLEINFO ABSTRACT Keywords: The Permian volcano-sedimentary strata of the Hida Gaien belt of SW Japan are characterized by Lower Permian Permian stratigraphy felsic tuffs and tuffaceous clastic rocks, Middle Permian andesites and sandstones rich in volcanic detritus, and Hida Gaien belt upper Middle to Upper Permian feldspathic sandstones and mudstones. A combination of petrographic and Sandstone provenance geochemical analyses together with zircon U-Pb dating for the sandstones reveals a provenance change from an Detrital zircons U-Pb dating immature undissected arc with active felsic and intermediate volcanism during the Early-Middle Permian to a Magmatic arc zircons with 500-400 Ma zircons. This age distribution suggests that sediment was derived from the Permian arc and early Paleozoic basement. Similar characteristics to Permian strata in the Hida Gaien belt are recognized in from volcanic to clastic rocks. In particular, Permian strata of the Jilin area (NE China) are lithostratigraphically similar to those of the Hida Gaien belt and have similar detrital zircon U-Pb ages. As such, Permian strata in both areas would have been deposited proximal to each other in the same tectonic setting along a single subduction zone in the western Paleo-Pacific Ocean. 1.Introduction particular region using these data alone. In addition, the complex dis- tribution of Permian terranes originating from accretionary wedges (the The Paleozoic terranes of Japan (Fig. 1) developed in an arc-trench Akiyoshi and Ultra-Tamba belts) and an oceanic island arc-back-arc system (or systems) along the active continental margin between the system (the Maizuru belt; e.g., Kojima et al., 2016) hampers a reasonable South China block and neighboring microcontinental blocks, such as the reconstruction that accounts for the tectonic setting of the entire proto- Jiamusi-Khanka and Bureya blocks of the eastern Central Asian Japan region during the Paleozoic. As such, determining the sedimen- Orogenic Belt (CAOB) (e.g., Isozaki et al., 2014; 2017; Isozaki, 2019). tary provenance of Upper Paleozoic strata helps us understand the tec- Representative terranes consisting of shallow-marine shelf deposits tonic setting of proto-Japan better because it provides information on include the Hida Gaien, South Kitakami, and Kurosegawa belts. These the geographic relationship of the sedimentary basins with neighboring belts are important for understanding the origin and early tectonic his- continents and island arcs and their behaviors as sources of clastic tory of “proto-Japan", which is characterized by Lower Paleozoic con- sediments (e.g., Yoshida et al., 1994; Yoshida and Machiyama, 2004; tinental （granitic) and ophiolitic basements and 1 overlying Takeuchi et al., 2008; Hara et al., 2018a; Zhang et al., 2018; Ohkawa Paleozoic-Mesozoic strata (e.g., Isozaki et al., 2015; Ehiro et al., 2016; et al., 2021). Previous studies of Permian brachiopod and fusulinid Ishiwtari et al., 2016). Faunal and floral data from the Middle-Upper faunas from the Hida Gaien and South Kitakami belts have revealed a Paleozoic strata of these terranes, however, are broadly similar to those paleobiogeographic correlation with the eastern CAOB in Northeast of the South China block and the CAOB (e.g., Ehiro and Kanisawa, 1999; China and Inner Mongolia (e.g., Shi and Tazawa, 2001; Shi, 2006; Ueno, Ueno, 2006; Kido and Sugiyama, 2011; Williams et al., 2014), and it is 2006). Therefore, the comparison among these belts and the eastern therefore difficult to determine a clear geologic relationship with a CAOB in terms of provenance and its temporal changes provides an * Corresponding author. E-mail address: nge02013.pls.foot22@gmail.com (K. Suzuki). https://doi.org/10.1016/j.jseaes.2021.104888 Available online 10 July 2021 1367-9120/@ 2021 Elsevier Ltd. All rights reserved. K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 opportunity for new insights to be gained into the tectonic setting of the basin on a continental shelf around theNorth China block and theCAOB eastern margin of the Asian continent. (e.g., Shi and Tazawa, 2001; Tazawa, 2001b). An alternative interpre- Prior to the opening of the Sea of Japan, the Hida Gaien belt, located tation is that the Paleozoic strata consist of various geological units at the northern edge of sW Japan, was directly connected to the areas of having faunal affinities with the South China block and the CAOB that Sikhote Alin and Northeast China (e.g., Kojima et al., 2000; Kemkin were juxtapositioned by Jurassic-Cretaceous dextral/sinistral strike-slip et al., 2016). According to the geotectonic division of Northeast Asia, movements (Ehiro et al., 2016; Tsukada, 2003). this belt corresponds to the southeastern termination of the CAOB and is In this study, we re-assess the stratigraphy, age, and provenance of thus pivotal in developing an understanding of the tectonic relationship the Permian strata of the Hida Gaien belt by using zircon U-Pb dating, between SW Japan and the Asian continental region. Several paleonto- sandstone petrography, and geochemistry. Zircon U-Pb dating of pre- logical studies have been conducted in the South Kitakami belt to viously undated strata in this belt allows a more robust understanding elucidate the belt's paleogeographic relationship with the Asian conti- of the belt's origin and provides an absolute temporal scale (e.g., nent during the Paleozoic (e.g., Ehiro, 1997; Tazawa, 2002); however, Suzuki et al., 2019). Sandstone petrography and geochemistry have the South Kitakami belt is isolated on the Pacific margin as a fault- been widely used to determine provenance and tectonic setting (e.g., bounded tectonic sliver, and its original relationship to the continental Bhatia, 1983; Dickinson et al., 1983; Bhatia and Crook, 1986; region remains uncertain. In this respect, the Hida Gaien belt, which was directly connected to the continent, provides an opportunity to explore effective when studying the evolution of an island arc through sedi- the relationship of the belt to the Asian continent. In the Hida Gaien belt, mentary provenance analysis (e.g., Greene et al., 2005; Cai et al., 2011; however, the stratigraphy and detailed ages of the Permian strata are Long et al., 2012; Hara et al., 2018a; 2018b). The present study re- poorly constrained, and only limited zircon dating has been conducted constructs the overall Permian stratigraphy of the Hida Gaien belt and (Nakama et al., 2010; Suzuki et al., 2019). Consequently, the late elucidates the sedimentary history and changes in provenance based Paleozoic history of the belt remains unclear. It has been suggested that on petrographic and geochemical analyses. We also discuss a new the Upper Paleozoic strata of this belt were deposited in a sedimentary interpretation of the tectonic relationship between the Hida Gaien belt Central Asian Orogenic Belt G Bureya block Sonid Zuoqi UCM QSuolun 45°N Songliao EHOB oxi Ujimqin Jilir NOB Basin Jiamusi-Khanka 1. 1n block SSZ/HB SOB Solonker bnoApluos Sea of Japan 39°N North China block SKB Fig. 2A distribution of Permian formations in Inner Mongolia and Northeast China HGB Qiling-Dabie Nagato Tectonicbelt 2 Pacific Ocean suture 33°N Kyushu KSB 0 500 km South China block 1149E F 130°E 139°E Selected Paleozoic to Mesozoic terranes of Japan Akiyoshi belt (Permian accretionay complex) Hida-Oki belt Maizuru belt (granite-gneiss complex metamorphosed at ca. 250 Ma) (Permian island arc and back-arc basin system) Hida Gaien, South Kitakami, and Kurosegawa belts Ultra-Tamba belt (Paleozoic arc-related rocks) (Permian accretionary complex) Renge metamorphic belt Mino-Tamba belt (Carboniferous high P/T metamorphic rocks) (Jurassic accretionary complex) Fig. 1. Index map of Japan, South and Northeast China, Inner Mongolia, and Far East Russia, showing the tectonic division of continental blocks, microcontinents, orogenic belts, sutures, and accretionary complexes (modified from Kojima et al., 2000; Takeuchi et al., 2008; Chen et al., 2017; Liu et al., 2017; Eizenhofer and Zhao, 2018; Wallis et al., 2020). SOB: southern early to mid-Paleozoic orogenic belt; SSZ/HB: Solonker suture zone / Hunshandake block; NOB: northern early to mid- Paleozoic orogenic belt; EHOB: Erenhot-Hegenshan ophiolite belt; HGB: Hida Gaien belt; SKB: South Kitakami belt; KSB: Kurosegawa belt. K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219(2021）104888 20.85 2.2.Fukuji area provides an updated view on the reconstruction of the Permian paleogeography of Northeast Asia. The Fukuji area is located 20 km east of the Moribu area (Fig. 2A). Paleozoic strata of the Fukuji area are subdivided into the Yoshiki, 2. Geologic outline 0.140 The Yoshiki Formation is characterized by felsic tuff and alternating The Hida Gaien belt is composed mainly of Ordovician to Lower tuffaceous sandstone and mudstone. Radiolarians, chitinozoans, and Cretaceous shallow-marine strata and is narrowly distributed be- U-Pb ages from this formation indicate a Late Silurian-Early Devonian tween the Hida-Oki (granite-gneiss complex metamorphosed at ca. age (e.g., Kurihara, 2004; Vandenbroucke et al., 2019). The Lower 250 Ma) and Mino (Jurassic accretionary complex) belts in central Devonian Fukuji Formation, which is partly coeval with the Yoshiki Japan. This belt extends to the Nagato Tectonic belt and central Formation, is composed of fossiliferous muddy limestone (e.g., Igo et al., Kyushu in the Inner Zone of Southwest Japan (e.g., Isozaki and 1975). The Ichinotani Formation is characterized by bedded limestone 0.263 partly interbedded with red mudstone (Igo, 1956). Fusulinids and small Shiroumadake areas in central Japan (Fig. 2A), a tectonic pile con- foraminiferas indicate that this formation is Early to Late Carboniferous sisting of the Hida Gaien belt and other Upper Paleozoic terranes (the 0.137 Renge metamorphic, Akiyoshi, and Maizuru belts) is preserved nabe, 1991). The Yoshiki, Fukuji, and Ichinotani formations are boun- (Komatsu, 1990) and is considered to reflect accretion in an active ded by faults continental margin during the Paleozoic. The Permian Mizuyagadani Formation is composed mainly of felsic Paleozoic strata of the Hida Gaien belt consist mainly of Ordovician tuff, mudstone, sandstone, and conglomerate. It has been subdivided 0.65 into lowermost, lower, middle, and upper members (Tsukada and clastic rocks, Devonian limestone, Carboniferous mafic and felsic vol- 0.020 canic rocks and limestone, and Permian intermediate volcanic and NE-SW and dip 60°-80° to the northwest (Fig. 3A). This formation is in clastic rocks (e.g., Tsukada et al., 2004; 2017; Kurihara, 2007; Kurihara faulted contact with the Ichinotani Formation, and their original strat- and Kametaka, 2008 and references therein). Permian strata are igraphic relationship might have been conformable or exhibited a minor 21.55 unconformity (Tsukada and Takahashi, 2000). The lowermost member and Ohno (Ohno City, Fukui Prefecture) areas in central Japan (Figs. 2 comprises calcareous clastic rocks and muddy limestone. The lower and 3). The following sections outline the stratigraphy and age of each CrO area, with particular focus on the Permian. mudstone (Niko et al., 1987; Kurihara and Kametaka, 2008). The radi- olarian assemblages suggest that the lower member is correlated with MgO 2.1. Moribu area doalbaillella lomentaria Range Zone (Asselian-Sakmarian) and the Paleozoic strata in the Moribu area (Fig. 2) are subdivided into the and upper members are composed of sandstone and conglomerate with Rosse, Arakigawa, and Moribu formations (e.g., Ehiro et al., 2016), mudstone. No age-diagnostic fossils have been reported from the middle 0.017 and upper members. 15.97 The Sorayama Formation conformably overlies the Mizuyagadani MnO Formation (Tsukada et al., 1999) and consists mainly of intermediate fossils (Leptophloeum), which assigns this formation to the Upper volcanic and volcaniclastic rocks with conglomerate and minor lime- Devonian (Tazawa et al., 2000). The Arakigawa Formation is charac- stone blocks (Tsukada and Takahashi, 2000). The conglomerate contains 0.022 limestone pebbles that contain abundant Late Carboniferous and Early fossils such as brachiopods, fusulinids, trilobites, and corals have been Permian fusulinids (Igo, 1956; Tsukada et al., 1999). Tsukada et al. reported from this formation (e.g., Suzuki and Shino, 2020; Tazawa (1999) also reported Middle Permian fusulinids, including Parafusulina et al., 2021 and references therein). 0.23 The Moribu Formation, originally defined by Isomi and Nozawa of the conglomerate. On the basis of these fusulinids, those authors (1957) in the southern part of this area (Fig. 2B), consists of Permian concluded that the Sorayama Formation is Middle Permian in age. clastic rocks with felsic tuff and limestone. The strata strike NE-SW to NNE-SSW and dip 30°-80° to the northwest. Yoshida and Tazawa 2.3.Ohno area (2000) subdivided the clastic rocks in this formation into basal, lower, middle, and upper members. Middle Permian brachiopods character- The Ohno area is located in the eastern part of Fukui Prefecture izing a mixed fauna of Boreal/bi-temperate/Tethyan species have been (Figs. 2A and 3B). The Paleozoic strata are subdivided into the Kagero, 0.15 Shibasudani, Kamianama, Nagano, Tomedoro, and Oguradani forma- lineatus, Leptodus nobills, Yakovlevia kaluzinensis, and Waagenoconcha cf. tions, which are bounded by faults (e.g., Kurihara, 2003). The Kagero imperfecta (Tazawa, 2001b). Monodiexodina sutchanica has also been Formation consists mainly of tuffaceous sandstone and mudstone, 0.835 including calcareous concretions from which Early Silurian radiolarians (2004) (see Ueno, 2006 for identification of the species). Limestone 0.15 pebbles within conglomerate of the basal member of the formation Formation is Upper S Silurian-Lower Devonian with abundant contain Early Permian fusulinids such as Pseudofusulina fusiformis and radiolarian-bearing horizons (Kurihara and Sashida, 2000a), and it is Misellina sp. (Yamada and Yamano, 1980). Poorly preserved Permian composed of sandstone and mudstone with minor limestone blocks. The radiolarians have been reported from felsic tuff in the underlying strata Kamianama Formation comprises muddy limestone that contains Early of the basal section (Umeda and Ezaki, 1997), but the fossil morphol- Devonian tabulate corals, trilobites, and cephalopods (Kurihara, 2003). ogies have not been sufficiently preserved to allow identification of the The Nagano Formation is characterized by massive limestone that con- tains middle Carboniferous fusulinids (Yamada, 1967). Although age- has reported detrital zircon U-Pb ages of 264-250 Ma for sandstones diagnostic fossils have not been reported from the Tomedoro Forma- 10.12 tion, it comprises mafic to intermediate volcanic rocks and is therefore considered to correlate with the Carboniferous Arakigawa Formation of to extend to the earliest Triassic. the Moribu area or with the mafic volcanic rocks of the Maizuru belt (e. 3 K.Suzuki and T.Kurihara Journal ofAsian Earth Sciences 219(2021）104888 136°13′E 137°E A 137°20E Renge Hirase distribution of Paleozoic formations metamorphic belt Kurabashira River in the Hida Gaien belt Maizuru belt 70 62N H1 Akiyoshi belt 264Ma Itoigar $2 37°N 36°16N H9 256Ma Sea of Japan Mt. Shiroumadake Douden HGB Kanazawa H6269 Ma 36°30N Moribu Fukuji (Fig. 2B) northern part (Fig. 3A) Hida belt Fukui Takayama Ohno Mino belt (Fig. 3B) 36°15N 50 km strike and dip of 137°18′35E bedding plane (tops unknown) Rosse-Kanayama (tops known) 46 R6 dated samples 270 Ma (see Fig. 4) southern part R1 282Ma fault 36°14N inferred fault 263Ma M13 38个 Quaternary pyroclastic rocks 262Ma Oamamiyama Group M5 (Upper Cretaceous) Moribudani River 272 Ma Tetori Group M12 1 km Moribu (Lower Cretaceous) Paleozoic strata of theMoribu area Douden Formation Moribu Formation (Permian) Arakigawa Formation (Jurassic) MB3 MB4 MB5 (Carboniferous) Funatsu granite Rosse Formation MB1 MB2 (Permian-Jurassic) (Devonian) Fig. 2. (A) Distribution of Paleozoic formations in the Hida Gaien belt with adjacent terranes, and (B) geologic map of the Moribu area. The zircon U-Pb ages of samples H1 and H6 are recalculated using the data of samples 0610-05 and 0529-63 reported by Suzuki et al. (2019), respectively. K.Suzuki and T.Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 (A) 137%30'38E 1379 16E B 36°13'35N iValley 35°52'25N Ichinotan 279'Ma MZ5 Ra3 Ra2 283 Ma MZ1 Ra1 Wasadani Valley Mizuyagadani Valley NEL.98 Haamidani 65 35°51N 500m 500m 136°3914E strike and dip of dated samples MZ1 Quaternary bedding plane (see Fig.5A) serpentinite pyroclastic rocks (tops unknown) fault accretionary complex of the Mino belt (Jurassic) cystalline schist (tops known) inferred fault Paleozoic strata of the Fukuji area Paleozoic-Mesozoic strata of the Ohno area Sorayama Formation Fukuji Formation Tetori Group Oguradani Formation (Middle Permian) (Lower Devonian) (Lower Cretaceous) (Middle Permian) Mizuyagadani Formation Yoshiki Formation Ohtani Formation Tomedoro Formation (Lower-Middle Permian) (Silurian-Devonian) (Triassic) (Carboniferous-Permian ?) Ichinotani Formation undivided strata MotodoFormation Kamianama Formation (Carboniferous) (Silurian-Permian ?) (Upper Permian-Triassic) (Lower-Middle Devonian) Fig. 3. Geologic maps of the (A) Fukuji and (B) Ohno areas, modified from Kurihar netaka (2008) and Kurihara (2003), respectively. Ra1-Ra3 indicate Early Permian radiolarian localities of Sakmarian (Niko et al., 1987), Asselian-Sakmarian , and Kungurian (Kurihara and Kametaka, 2008), respectively. g., Tazawa and Matsumoto, 1998; Matsumoto, 2012). felsic tuff, tuffaceous sandstone and mudstone, and siliceous mudstone. The Permian Oguradani Formation is composed of alternating The horizon containing poorly preserved radiolarians reported by mudstone and sandstone beds with argillaceous limestone and calcar- Umeda and Ezaki (1997) is included in this unit. The felsic tuff is eous sandstone. These rocks strike WNW-ESE to ENE-WSW and dip yellowish white in color and is intercalated with black siliceous 40°-80° to the south but have been stratigraphically overturned mudstone (Fig. 6A). The tuff includes abundant volcanic glass shards (Fig. 3B). This formation has yielded Middle Permian brachiopods and with euhedral plagioclase (Fig. 6B) and is commonly associated with fusulinids (Tazawa and Matsumoto, 1998; Ueno and Tazawa, 2004), tuffaceous sandstone and mudstone measuring several centimeters which characterize the mixed fauna of Boreal/bi-temperate/Tethyan thick. The sandstones are fine- to medium-grained with graded bedding species similar to those of the Moribu Formation. and parallel and cross lamination, indicating a turbidite affinity. Unit MB2 consists of conglomerate, sandstone, and limestone. This 3. Lithostratigraphic descriptions of the Permian strata unit corresponds to the basal member described by Yoshida and Tazawa (2000). The conglomerates, which are commonly observed in the lower 3.1. Moribu Formation (Moribu area) part of the unit, contain subangular to rounded granules and pebbles of limestone and volcanic rocks (Fig. 6C). The sandstone in this unit is dark The Moribu Formation consists of felsic tuff, tuffaceous sandstone gray in color and is characterized by abundant fine- to medium-grained and mudstone, siliceous mudstone, mudstone, sandstone, and plagioclase grains and lithic fragments of andesite (Fig. 7A). The lime- conglomerate (Fig. 4). The total thicknessis is >1500 m. In this study, we stone is blueish gray in color and contains ooids and abundant bioclasts, subdivide this formation into units MB1, MB2, MB3, MB4, and MB5 in such as fusulinids and crinoid stems, which are filled with calcareous ascending stratigraphic order, on the basis of their lithologic features mudstone and minor calcite cement. (Fig. 4). The characteristics of the sandstones are described in detail in Unit MB3 is composed of alternating sandstone and mudstone Section 5.1. (Fig. 6D) and a small amount of limestone. In the sandstone-mudstone Unit MB1, exposed at Rosse-Kanayama (Fig. 2B), is characterized by succession, the sandstone grades gradually to mudstone in each bed, and K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021） 104888 Moribu Formation (Permian) (10) H10 oer 256 Ma H9 Upp (4） H8 M13 263 Ma MB5 200 m (1) (3) northern part M11 southern part (9) -M10 269 Ma -M9 (8) H6 MB4 -M8 H4 (2) M12272Ma pp Mis -H5. P M6 M7 (7) MB3 H2 262 Ma 1 264 Ma M5 270 Ma M4 (5)R6 sample horizon and location R5 M3 (6) R4 MB2 M1,2 R3- M13 dated samples (see Fig. 2B) R2. MB1 282 Ma R1 stratigraphic upper (normal grading) southernpart northern part Rosse- (6) Kanayama lithostratigraphic boundary (7) R5 H3- IR4(5) R6 H1 nudstone (10) andstone H4 conglomerate R1 (4) Hirase R2 H8- H9 alternating sandstone M132 M8 M3 (8) imestone H10- and mudstone M10 (1)) M4 fM1 H2 M2 Douden felsic tuff, tuffaceous H5 M6 H6 sandstone, tuffaceous mudstone Moribu ((2) M7 (9) Fig. 4. Stratigraphic columns of the Permian Moribu Formation in the Moribu area showing sample horizons and locations. Samples shown in the figure were used S p ss po se q n a o u a s n a ss a xa ss l p od the mudstone shows parallel laminations of fine-grained sandstone. and is locally interbedded within the alternating sandstone and Some sandstones contain abundant volcanic lithic fragments, which are mudstone. similar to those of unit MB2, with pebble- to cobble-sized fragments of white in color and has been heavily recrystallized. 3.2. Mizuyagadani and Sorayama formations (Fukuji area) Unit MB4 consists mainly of massive and thickly bedded sandstones with mudstone intercalations. The sandstone is gray in color and fine to The Mizuyagadani Formation consists of calcareous mudstone, muddy limestone, felsic tuff, tuffaceous sandstone and mudstone, sili medium grained. The mudstone intercalations are black in color and several millimeters to less than 10 cm in thickness. The mudstone shows ceous mudstone, mudstone, sandstone, and conglomerate (Fig. 5A). The total thickness is ~350 m. Stratigraphic subdivisions and lithologic de- parallel laminations of thinly bedded, very fine-grained, brownish-gray sandstone measuring several millimeters thick. scriptions largely follow those of Kurihara and Kametaka (2008), with Unit MB5 is mudstone dominated and is partly interbedded with further information based on our reinvestigation. alternating sandstone and mudstone and thin tuffaceous sandstone The lowermost part of the Mizuyagadani Formation is composed of layers. The mudstone is blueish gray in color. The mudstone is inter- calcareous sandstone and mudstone, as well as muddy limestone. These bedded by sandstone layers that grade from sand- to mud-sized grains rocks are black to dark gray in color and well bedded with beds and exhibit parallel lamination (Fig. 6F). The sandstones are gray in from several millimeters to centimeters in length, including crinoid stems, bryozoans, gastropods, and brachiopods. The lower part of the formation consists mainly of radiolarian-bearing felsic tuf and K. Suzuki and T. Kurihara 0.314 (B) A SR2 Sorayama Fm. (1) (1) -SR1 (Middle Permian) (1) (1) 279 Ma rmian) -OG2 MZ5 MZ5 OG1-! MZ4 upper part MZ3 0G3 Middle Per MZ4 Permian) MZ2 MZ1 Wasadani Valley Mizuyagadani Valley part (Middle I Oguradani middle part Formation -MZ3 middle I -MZ2 1100 m OG1 MZ1 100 m lower part 283 Ma lowermost 10 M 0 part sample horizon and location mudstone sandstone conglomerate MZ1 dated samples (see Fig. 3A) calcareous alternating sandstone limestone clastic rocks Iand mudstone 个 35.63 felsic tuff, tuffaceous andesite lithostratigraphic boundary sandstone, tuffaceous mudstone Fig. 5. Stratigraphic columns of Permian formations in the (A) Fukuji and (B) Ohno areas showing sample horizons and locations. The stratigraphy for the Miz- uyagadani and Sorayama formations is from Kurihara and Kametaka (2008). Samples shown in the figure were used for petrographic and geochemical analyses, except for sample MZ1. alternating tuffaceous sandstone and mudstone. The felsic tuff is The Sorayama Formation consists mainly of andesitic volcanic rocks yellowish white and pale green in color and is commonly intercalated and volcaniclastic rocks (Fig. 5A). The conformable contact between the with gray sandstone and black mudstone (Fig. 6G). The mudstones in Mizuyagadani and overlying Sorayama formations has been described this part contain thin felsic tuff and sandstone layers that grade from by Tsukada et al. (1999). The andesitic volcanic rocks of this formation sandstone- to mudstone-dominated beds. Some of the mudstones occur as lavas and dikes. They are highly altered, are blueish gray in contain abundant sponge spicules and small numbers of radiolarian 0.605 fossils (Kurihara and Kametaka, 2008). Each bed of tuff, sandstone, and have been replaced by chlorite or calcite with opaque minerals. The mudstone varies in thickness from several millimeters to several tens of groundmass includes abundant acicular plagioclase with intersertal centimeters. The sandstones tend to be relatively thick. The middle part texture (Fig. 61). The volcaniclastic rocks commonly occur as pebble- to of the formation is composed of sandstone, mudstone, and conglom- cobble-sized breccia containing andesite and tuff. These rocks are erate, with alternating layers of these lithologies. The sandstone is dark massive to weakly stratified and are black, dark green, or reddish brown gray in color, is medium to coarse grained, and contains local thinly in color. Conglomerate beds commonly occur within the breccia as bedded mudstones with parallel and cross laminations. Each sandstone interbedded layers of several meters to several tens of meters thickness 0.743 17.03 ceous. The conglomerate contains subangular to subrounded granules felsic-intermediate tuff, basalt, and porphyrite. The clasts are poorly and pebbles of limestone and volcanic rocks. Some calcareous bioclasts, sorted, subangular to rounded, and pebble to cobble size. The con- such as fusulinids, are also contained in the matrix. The upper part of the glomerates are matrix supported or partly clast supported, and they have formation consists mainly of alternating sandstone and mudstone. The a coarse-grained, dark green, tuffaceous sandstone matrix. Some con- 0.800 glomerates contain abundant limestone clasts in a tuffaceous matrix. to abundant volcanic lithic fragments. Limestone clasts are 15-20 cm in diameter and subangular to rounded in K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 Bt E Cy 3cm ms 500μm Fig. 6. Photographs and photomicrographs showing the main lithologies of the (A-F) Moribu, (G-H) Mizuyagadani, and (I) Sorayama formations (see Figs. 4 and 5 ( n a yo () d-) () n o d ) n go n n () (ss os Alternating sandstone and mudstone of unit MB3. (E) Polished surface of sandstone in (D), showing clasts of volcanic rocks. (F) Alternating sandstone and mudstone of unit MB5. (G) Felsic tuff of the lower part of the Mizuyagadani Formation. (H) Sandstone of the upper part of the Mizuyagadani Formation. (I) Photomicrograph of andesite (cross-polarized light). ms: mudstone; ss: sandstone; Pl: plagioclase; Bt: biotite; vgs: volcanic glass shard; Cv: clast of volcanic rocks. shape, and the limestone-dominated conglomerates are matrix sup- The upper part of the formation is composed of limestone, calcareous sandstone, and slaty mudstone. The limestone and calcareous sandstone tuffaceous and calcareous sandstone and contains Middle Permian fu- are exposed in the Oguradani Valley, where the type locality of this sulinids (Tsukada et al., 1999). Felsic tuff is also observed in some lo- formation is defined. The limestone is black in color, is argillaceous, and contains abundant crinoid and brachiopod fossils (Tazawa and Matsu- and contains local thin limestone intercalations. moto, 1998). The slaty mudstone is black to dark blue in color and contains thin intercalations of fine-grained sandstone. 3.2.1. Oguradani Formation (Ohno area) The Oguradani Formation is composed of alternating mudstone and 4. Analytical methods sandstone beds, with argillaceous limestone and calcareous sandstone (Fig. 5B). The total thickness is >600 m. This formation is subdivided 4.1. Sandstone petrography into lower, middle, and upper parts (Tazawa and Matsumoto, 1998). The lower part of this formation is composed of mudstone and A total of 35 medium-grained sandstone samples were collected from sandstone. The mudstones and sandstones are typically light gray to the Moribu, Mizuyagadani, and Oguradani formations for modal anal. white in color, with distinctive shearing common throughout (Otoh ysis (Figs. 4, 5, and 7). For each thin-section, 500 detrital grains were et al., 2004). The mudstones are partly pale green in color and siliceous. counted using the Gazzi-Dickinson method (Ingersoll et al., 1984). The sandstones contain rare, poorly-preserved brachiopods and bryo- When the crosshairs on the microscope intersected a lithic fragment zoan fossils. Monodiexodina is observed in the sandstone (Ueno and derived from granitic rocks, the grain was counted as quartz or feldspar Tazawa, 2004). rather than a lithic fragment. All counted parameters are listed in Sup- The middle part of the formation is dominated by alternating beds of plementary Table 1. sandstone and mudstone, which are gray and black in color, respec- 4.2. Whole-rock geochemistry The sandstone exhibits graded bedding from medium to fine grained. Thickly bedded medium- to coarse-grained sandstones also occur. The The 35 samples collected for sandstone petrographic analysis were limestone is massive, white to gray in color, and strongly recrystallized. also used for whole-rock geochemical analysis. Two additional samples No fossils other than crinoids have been observed in this limestone. of volcanic rocks were collected from the andesitic dike (SR1) and lava K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 500 μum 500μum 00pm Fig. 7. Photomicrographs of representative sandstones from the (A-E) Moribu and (F) Mizuyagadani formations (see Figs. 4 and 5 for sample horizons and locations). (A) Lithic wacke of unit MB2 (sample M1). (B) Lithic arenite of unit MB3 (sample M5). (C) Lithic wacke of unit MB3 (sample R4). (D) Feldspathic arenite of unit MB4 (sample M9). (E) Feldspathic wacke of unit MB5 (sample H10). (F) Lithic arenite of the upper part of the Mizuyagadani Formation (sample MZ4). Qz: quartz; Pl: plagioclase; Kf: K-feldspar; Lv: volcanic lithic fragment; Lp: plutonic lithic fragment; Lm: metamorphic lithic fragment. (SR2) in the Sorayama Formation (see Figs. 5A and 61) for analysis. The 3A, 4 and 5A). Detrital zircon U-Pb ages from three sandstone samples 37 samples were crushed to powder using a tungsten carbide mortar and of units MB3-MB5 of the Moribu Formation (H1, H6, and H9; Figs. 2B agate ball mill. The contents of 10 major and 11 trace (Ba, Cr, Nb, Ni, Rb, and 4) have already been reported in Suzuki et al. (2019). Among these Sr, V, Y, Zr, Pb, and Th) elements were determined by X-ray fluorescence samples, sample H9 has been analyzed again with additionally separated (XRF) spectrometry (Rigaku RIX3000) on glass beads at Nigata Uni- zircon grains in the present study to verify the reproducibility of versity, Japan (Table 1). The total Fe content is reported as Fe2O3. Loss 500-400 Ma zircon data used in the discussion (Section 6). For samples on ignition (LOl) values were measured by weighing the samples before H1 and H6, we recalculated the data from Suzuki et al. (2019) and re- and after 7h of heating at 900 °C. Analytical procedures and accuracy assessed their depositional ages. Table 2 and Supplementary Table 3 have been described by Takahashi and Shuto (1997). Contents of Sc and show the zircon U-Pb ages for all ten samples (R1, M5, R6, M12, M13, H1, H6, H9, MZ1, and MZ5). ma-mass spectrometry (ICP-MS) using an Agilent 7500a instrument at Zircon grains were manually separated from the samples using Nigata University. Liquid samples for analysis by ICP-MS were pre- magnetic and methylene iodide heavy liquid separation techniques. pared based on a combined acid digestion procedure (HCl, HF, and Zircon grains were picked at random under a microscope and embedded HNO3). The analytical precision of the Sc and REE contents, as estimated in epoxy resin (Petropoxy 154). The surface of embedded zircon grains by relative errors from the literature values of the geological reference was polished using 1 μm diamond paste on a cloth. Internal structures materials BHVO-2 and W-2 (U.S. Geological Survey) and JB-2 and zoning within the polished zircon grains were observed using a ia ( ) scanning electron microscope (SEM, JEOL JSM-5600) equipped with a 8%. Analytical accuracy, estimated by relative standard deviations in Cathodoluminescence imaging system (Gatan Mini CL) at Nigata Uni- replicated analyses of BHVO-2 and W-2, was better than 6%. All figures versity to select suitable spots for analysis. showing data, such as Harker variation and discrimination diagrams, are presented using anhydrous values of the major-element oxides normal- ized without LOI. 7500a quadrupole ICP-MS instrument coupled to a New Wave UP213 LA system at Nigata University. We followed the methods of Ueda et al. (2018). Sample aerosols were generated from ablation pits (30 μm 4.3. Zircon U-Pb dating diameter) using a laser frequency of 5 Hz (213 nm wavelength) at an energy density of ~10 J/cm2. Total analysis durations were 28 s during Two felsic tuff and eight sandstone samples were collected from the an acquisition time (60 s), with the resultant isotope ratios being the Moribu and Mizuyagadani formations for zircon U-Pb dating (Figs. 2B, K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 Table 1 Averages of whole-rock major-element oxides (anhydrous normalized basis) and selected trace-element and REE compositions from sandstones and volcanic rocks of the Hida Gaien belt. Area Moribu Fukuji Ohno Formation Moribu Mizuyagadani Sorayama Oguradani Lithology sandstone sandstone andesite sandstone Lithostratigraphic subdivision MB2 MB3 MB4 MB5 middle-upper parts -2 middle part Number of analyzed samples 2 12 10 3 4 4 Major element oxides (wt.%) SiO2 57.64 66.90 67.42 70.73 63.94 57.34 66.80 TiO2 0.97 0.79 0.72 0.62 0.95 0.96 0.76 Al2O3 18.14 16.74 15.12 14.82 14.43 16.89 16.18 Fe2O3 8.24 5.52 5.07 3.99 6.50 7.12 6.46 MnO 0.11 0.08 0.10 0.06 0.16 0.12 0.11 MgO 4.92 1.90 2.01 1.55 1.75 4.47 2.35 CaO 5.20 2.60 3.44 3.14 6.92 8.62 1.56 Na20 3.44 3.63 4.58 3.47 4.42 2.73 4.38 K20 1.13 1.74 1.44 1.46 0.77 1.48 1.27 P2O5 0.21 0.09 0.10 0.14 0.17 0.28 0.12 Total 100.00 100.00 100.00 100.00 100.00 100.00 100.00 earthelements(ppm) Ba 402.95 412.94 389.34 255.20 125.85 384.75 240.73 1D 44.30 26.93 23.92 28.40 22.95 179.30 30.23 Nb 8.09 10.54 7.71 8.88 4.08 6.63 3.97 Ni 20.75 7.43 5.75 11.53 6.08 32.85 7.28 Rb 18.37 42.80 34.00 43.94 18.18 41.38 31.94 Sr 905.29 406.30 386.98 501.36 313.24 428.82 333.90 V 175.64 101.25 95.63 66.68 112.88 148.97 124.13 Y 18.13 22.77 21.63 19.79 29.84 21.98 23.58 Zr 122.80 169.77 150.28 167.22 117.94 154.91 148.34 Pb 5.84 9.82 8.82 11.70 6.74 7.85 7.58 Th 4.23 7.75 7.40 8.03 3.40 3.37 4.95 Sc 17.22 11.20 10.65 6.99 14.31 18.91 11.25 17.38 La 13.95 15.07 19.56 10.50 17.31 12.35 mean of 20 s of time-resolved analysis after removing the first 8 s of data. Density Plotter (YCM2o; Vermeesch, 2012) has been widely used for a Instrumental mass bias and drift were corrected on the basis of eight maximum depositional age of a sample (e.g., Endo et al., 2018; Nagata repeated measurements of the NIST SRM610 silicate glass (Walder et al., et al., 2019; Miyazaki et al., 2019; Tokiwa et al., 2021). As mentioned 1993) per a cycle. Downhole Pb/U fractionation was corrected by below, the MsWD values of all samples in the present study were more (238U_206pb analysis of the 91500 zircon standard than 3.3 (Table 2). Therefore, the YCM2o ages were used to determine maximum depositional ages of all samples of the Moribu (R1, M5, R6, were obtained indirectly by multiplying the measured 207pb/206pb, M12, M13, H1, H6, and H9) and Mizuyagadani (MZ1 and MZ5) for- 206pb/238U, and natural 238u/235u (137.88) ratios. Data accuracy was mations. Results are presented in Table 2, together with the weighted monitored using analyses of the Plesovice zircon standard (238u_206pb mean age and youngest cluster age calculated by Isoplot v. 4.15 (Lud- age =337.13 ±0.37 Ma; Slama et al., 2008). The weighted mean wig, 2012) for each sample. Data for samples H1 and H6 were recal- culated using supplementary tables of Suzuki et al. (2019). Data for 335.6± 5.8 to 330.7±4.2 Ma during the analysis period and are sample H9 are the combined results of this study and Suzuki et al. consistent with the reference data within 2% error (Supplementary (2019). Table 3). In all data, isotopic errors on 206pb/238U are less than 5%. A correction for common Pb was not applied to the data. The evaluation of 5. Results the concordance of the dated zircon grains was performed using the method of Suzuki et al. (2019). All errors are stated with 2 standard 5.1. Sandstone petrography error (2o) analytical uncertainties. All ages and probability plots in Fig. 13 were calculated and plotted using Density Plotter (Vermeesch, As described above, most of the Permian formations of the Hida 2012). Detrital zircon U-Pb age spectra and probability plots in Fig. 15 Gaien belt contain abundant sandstones (Figs. 4 and 5), other than the were calculated using Isoplot v. 4.15 (Ludwig, 2012). Discordant data tuffaceous lithology of the Moribu Formation (unit MB1), the lowermost were rejected on the basis of anomalous isotopic data, and data that to lower parts of the Mizuyagadani Formation, and the Sorayama For- were collected too close to inclusions and cracks were discarded. For this mation. To investigate the provenance of the sandstones, we conducted study, only concordant data were used to construct probability plots modal analysis of the medium-grained sandstones obtained from units (Figs. 13 and 15). MB2 to MB5 of the Moribu Formation, the middle and upper parts of the For evaluating the depositional age of a sample, we calculated the Mizuyagadani Formation, and the middle part of the Oguradani For- weighted mean age of the youngest population overlapped with 20 mation. The main components of the sandstone are monocrystalline uncertainties of dated zircon grains and its mean square weighted de- quartz, plagioclase, K-feldspar, and volcanic, sedimentary, and meta- viation (MSWD) value. In general, if the MSWD value of the analyzed morphic lithic fragments (see Supplementary Table 1). population is closer to 1, the data used to produce the weighted average Sandstones of the Moribu Formation can be classified as the is considered to be near to a single population. However, if it is not following two types. The first is lithic arenite-wacke, which includes determined to be a single population, the youngest cluster age can be abundant fragments of mafic to intermediate volcanic rocks that are used to constrain a maximum depositional age for clastic rocks. consistent with those observed in units MB2 and MB3 (Fig. 7A, B, and C). Recently, the youngest cluster age calculated by Mixture Models of 10 K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 sorted and angular to subangular in shape, and are bound by muddy or tuffaceous matrices. Volcanic lithic fragments are characterized by lath- sandstor shaped plagioclase crystals. Plagioclase observed as monocrystalline 088 3 detritus is partly replaced by saussurite filled with phlogopite. The minor constituent grains are quartz, calcite, biotite, zircon, chlorite, and Fukuji opaque minerals. Around 40%, a range on average, of the total number 283±2 (100%) of grains counted per sample are volcanic lithic fragments. The per- felsic tuff 297-269 centage of lithic fragments (volcanic, sedimentary, and metamorphic) of 2719 the total grains ranges from 36 to 76%, with an average of 58%. The contents of monocrystalline quartz and feldspar are relatively low, with the lowest monocrystalline quartz content (<1%) being recorded in unit s s (%6千12) (18-44%). The second sandstone type in the Moribu Formation is feldspathic 9 arenite-wacke, which contains abundant monocrystalline quartz and feldspar with fewer lithic fragments (Fig. 7D and E). This sandstone type commonly occurs in units MB4 and MB5. Detrital grains of quartz and feldspar are moderately sorted and subangular to subrounded in shape, (%01 0) in a very fine-grained sandy to muddy matrix. The monocrystalline 646 quartz commonly shows undulatory extinction. Myrmekite is commonly 30 observed in feldspar grains. Volcanic lithic fragments are mainly felsic volcanic rocks such as rhyolite and dacite and differ markedly from those of units MB2 and MB3 that are characterized by the abundance of mafic to intermediate volcanic lithic fragments. The minor constituent 256 grains are biotite, muscovite, tourmaline, hornblende, sphene, zircon, 16.55 2.3 and opaque minerals. Monocrystalline quartz and feldspar contents sand average at 21% (14-30%) and 52% (49-64%), respectively. In unit MB3 of the Moribu Formation, several sandstones are inter- (26 12) sandstones contain not only fragments of mafic to intermediate volcanic 95 rocks but also substantial amounts of quartz and feldspar (~50%). 306 Taken together, the sandstones of the Moribu Formation show an overall change from lithic sandstones containing abundant mafic-intermediate Moribu volcanic lithic fragments to feldspathic sandstones, with sandstones showing intermediate characteristics of these two types occurring in the transition between the two. Sandstones in the middle to upper parts of the Mizuyagadani For- mation are characterized by abundant fragments of mafic to interme- diate volcanic rocks, which make up over half of the total grain populations (Fig. 7F) and resemble the lithologies in units MB2 and MB3 of the Moribu Formation. The detrital grains are poorly sorted and -258 97 subangular to subrounded in shape in a muddy and tuffaceous matrix with some development of calcite cement. Volcanic lithic fragments have intersertal texture with lath-shaped plagioclase crystals. The minor constituent grains are biotite, zircon, chlorite, and opaque minerals. Total lithic fragment contents average at 62% (54-78%), which is (512 182) roughly consistent with the modal data for the sandstones of units MB2 6好 and MB3. In the middle part of the Oguradani Formation, the sandstones are belt lithic arenite and contain substantial numbers of feldspar grains. The detrital grains are moderately sorted and subangular to subrounded in 282±1 (100%) shape and are set in a very fine-grained sandy to muddy matrix. The tuff 308-271 lithic fragments comprise mainly andesite, rhyolite, and felsic tuff. MR felsic 59 Myrmekite and perthite are commonly observed in feldspar grains. The minor constituent grains are hornblende, calcite, zircon, biotite, and opaque minerals. Average feldspar and total lithic fragment contents are 37% (32-41%) and 43% (41-45%), respectively. The sandstones of the tigraphic subdivision Oguradani Formation are similar to those in unit MB3 of the Moribu zircon U-Pb Formation in terms of their compositions, with both including abundant rdant volcanic lithic fragments and feldspar grains. (Ma) When plotted in a Qm-F-Lt diagram (Dickinson et al., 1983), data from units MB2 to MB5 of the Moribu Formation plot broadly from the tion ogy and total lithic fragment contents increase and decrease upward through the stratigraphy, respectively, although they commonly vary within individual lithostratigraphic units. Data for the Mizuyagadani 11 K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 Qm 35.63 1.72 35.53 14.80 The unit MB3 samples have more variation in the Eu anomalies, LREEs, and HREEs (Gd-Lu) than observed in other units. For example, in sample M4 of unit MB3, a flat REE pattern is recognized, showing the largest depletion in LREEs among all samples of this unit and a slight positive Eu MnO samples are generally similar to that of UCC, although the LREEs are Mixed slightly depleted. In addition, the unit MB3 samples except for sample M4 have Eu anomaly values from 1.02 to 0.70, and some of the samples show negative Eu anomalies. Units MB4 and MB5 show slight depletion in LREEs relative to UCC, and have negative Eu anomalies and flatted 10.22 units MB4 and MB5 samples are less than 0.97, trending negatively. Sample H3 from unit MB4 yields the lowest Eu anomaly value of 0.68. 22.19 Transitional arc 0.47 UDA FeO* Moribu Formation Lt 2.09 Oguradani Formation samples, in that LREEs are distinctively depleted relative to UCC values. MB4 ●MB5 However, the Eu and HREE values of same samples from the Oguradani MB2 MB3 Mizuyagadani Formation Formation differ markedly from those of UCC. As mentioned above, this Fig. 8. Qm-F-Lt diagram showing the tectonic fields of Dickinson et al. (1983). 21.25 Sandstone samples are from the Moribu, Mizuyagadani, and Oguradani for- Formation. 17.72 Chondrite-normalized REE plots for andesite samples from the Sor- total lithic fragments (lithic fragments + polycrystalline quartz); UDA: undis- 10.63 sected arc. 0.51 Al2O3 94.74 and largely overlap the plot for unit MB3 of the Moribu Formation contrast to that of UCC. Among the sandstone samples, the REE contents 36.08 of samples from unit MB2 of the Moribu Formation are particularly tional arc” sector, which overlap the plots for units MB3 and MB4 similar with those of the Sorayama Formation in terms of their REE (Fig. 8). patterns and lack of negative Eu anomalies (average 1.09). On Al2O3/SiO2- and TiO2-(Fe2O3 + MgO) diagrams (Bhatia, 1983), 5.2. Whole-rock geochemistry sandstone compositions plot broadly from the oceanic island arc to active continental margin fields (Fig. 11A and B). Data from unit MB2 of 95.28 13.60 major-element oxides from sandstone samples of the Moribu, Miz- For unit MB3, the compositions plot in a broad range from the oceanic uyagadani, and Oguradani formations and andesite samples of the Sor- MgO u sn n p ae n n MB5 tend to plot in the continental island arc and active continental 1.89 margin fields. Data from the Mizuyagadani and Oguradani formations 0.01 plot in the intermediate domain between the oceanic island arc and axes and contents of the measured elements as the vertical axes are continental island arc fields. presented in Fig. 9. All REE data obtained in the present study are pre- 35.42 sented as chondrite-normalized REE plots (Fig. 10) and are compared 5.97 with upper continental crust (UCC) values (Rudnick and Gao, 2014). (Fig. llC-E). Contents of V and Sc are useful indicators for assessing SiO2 contents in the sandstone samples vary from 51 to 75 wt% changes in mafic volcanic contribution to clastic rocks (V-Sc diagram, (Table 1). Sandstone samples from unit MB2 of the Moribu Formation Fig. 11C; Ryan and Williams, 2007; Surpless, 2015). Thorium is a typical 22.46 incompatible element, and zirconium is a recycled element, both of ples exhibit a trend of increasing average SiO2 contents upward through which are useful for estimating sediment supply from felsic volcanic 1.17 21.50 3.90 93.79 have intermediate Si02 contents (average of 64 wt% and 67 wt%, sedimentary basins (Bhatia and Crook, 1986). Variations in provenance respectively) that are similar to those from units MB3 and MB4 of the and tectonic setting can be constrained using the mean contents of these Moribu Formation. Harker variation diagrams (Fig. 9) show negative 0.80 correlations between SiO2 and Al2O3, TiO2, MgO, Fe2O3, and V. volcanic rocks (F), and Phanerozoic granite (G) (Condie, 1993; Fig. 11D Although there is variability among the samples from unit MB3 of the and E). Moribu Formation, positive correlations are observed between SiO2 and 37.08 K2O, Zr, and Th. Na2O shows no correlation with SiO2 content in any of sandstone samples from the Moribu Formation decrease roughly in the samples. The volcanic rocks of the Sorayama Formation (SR1 and ascending stratigraphic order from unit MB2 to unit MB 5. Diagrams of 6.02 Th/Sc-Zr/Sc (McLennan et al., 1993) and La-Th-Sc (Bhatia and Crook, 18.79 0.07 16.25 in a range from the Paleozoic andesite to Paleozoic felsic volcanic rock (1989). 25.48 12 K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 20 Al2O TiO NazO 19 18 oxides wt.% 17 1.0 16 15 ).6 14 13 0.4 50 55 60 65 70 75 80 50 55 60 65 70 75 80 50 55 60 65 70 75 80 10 K2O MgO Fe2O3 wt.% Oxides 50 55 60 65 70 75 80 50 55 60 65 70 80 50 5 50 6 70 75 80 300 12 200 Zr 11 Th 180 250 10 160 140 200 120 wda 100 150 > 80 60 100 3 20 50 55 60 65 70 75 80 50 55 60 65 70 75 80 50 55 60 65 70 75 80 SiO2 wt.% SiO2 wt.% SiO2 wt.% sandstone Moribu Formation andesite MB4 OMB5 Oguradani Formation Sorayama Formation MB3 Mizuyagadani Formation MB2 Fig. 9. Harker variation diagrams for selected major-element oxides (anhydrous-normalized values) and trace-element compositions of sandstones from the Moribu, Mizuyagadani, and Oguradani formations and andesites from the Sorayama Formation. SiO2 content is used for the horizontal axis. across the oceanic island arc and continental island arc fields (Fig. 11E). sandstone are euhedral in shape and 50-250 μm in diameter. Almost all In particular, data for unit MB2 of the Moribu Formation and the Miz- of the zircons show clear oscillatory zoning in CL images (Fig. 12), uyagadani Formation plot mainly in and around the oceanic island arc indicating that the zircons crystallized during a single period of field. Data for unit MB3 vary from near the oceanic island arc field to magmatic activity (Corfu et al., 2003; Hoskin and Schaltegger, 2003). within the continental island arc field. Data for units MB4 and MB5 plot For such magmatic zircons obtained from tuffs and clastic rocks, the largely in the continental island arc field. Compositions of the Oguradani weighted mean age with 2o uncertainty for single populations is useful Formation plot across the oceanic island arc and continental island arc fields. Spencer et al., 2016). In the present study, the MSWD values of all samples were more than 3.3 (Table 2); the data used to produce the 5.3. Zircon U-Pb dating weighted average is not considered a single population. Therefore, as mentioned in Section 4.3, the present study used the YCM2o ages Zircon U-Pb ages of the Moribu and Mizuyagadani formations are calculated by Mixture Models of Density Plotter (Vermeesch, 2012) to determine the maximum depositional ages of all samples. listed in Supplementary Table 3. Zircons obtained from felsic tuff and K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 (A) sandstone 1000 1000 Moribu Formation Eu/Eu*=1.09 Moribu Formation Eu/Eu*=0.92 /CI-chondrite /CI-chondrite UCC UCC MB2 MB3 100 100 Sample/( 10 10 1.0 1.0L La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1000 1000 Moribu Formation Eu/Eu*=0.87 Moribu Formation Eu/Eu* = 0.86 I-chondrite UCC UCC MB4 -chondr MB5 100 100 Sample/ 10 Sampl 10 1.0 1.0L La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1000 1000 Eu/Eu* = 0.94 Eu/Eu* = 0.87 Mizuyagadani Formation Oguradani Formation I-chondrite I-chondrite UCC UCC 100 100 Sample/CI 10 10 1.0 1.0 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu (B) andesite 1000 Sorayama Formation Eu/Eu* = 0.99 cite UCC I-chondr 100 Sample/ 10 1.0 Aa q g s Fig. 10. Chondrite-normalized REE plots of (A) sandstones from the Moribu, Mizuyagadani, and Oguradani formations, and (B) andesites from the Sorayama Formation. Chondrite-normalizing factors are based on Barrat et al. (2012). UCC: upper continental crust (Rudnick and Gao, 2014). unit MB5 yields a YCM2o age of 256 ±1 Ma. This sample contains and sandstone samples of the Moribu Formation have a broad range 562-378Ma zircons; these latest Neoproterozoic to Middle Paleozoic from 2353 to 241 Ma, although Proterozoic zircons are rare (Fig. 13A zircons are not observed in other samples from the Moribu Formation (Fig. 13A). Moribu Formation yields a prominent age peak at 308-271 Ma; a YCM2 Data for tuff sample MZ1 from the lower part of the Mizuyagadani age of this grain population is 282 ± 1 Ma. The rest of the samples from Formation are distributed in a single peak at 297-269 Ma, with a single the Moribu Formation have narrow age peaks within 335-250 Ma. Data 403 Ma zircon grain (Fig. 13B and Table 2). Zircons from this peak for sandstone samples M5, R6, and H1 from unit MB3 yield YCM2o ages population yield a YCM2o age of 283 ± 2 Ma. Data from sandstone of 262±2 Ma, 270 ±1 Ma, and 264 ±1 Ma, respectively. Sandstone sample MZ5 from the upper part of the Mizuyagadani Formation yield a samples M12 and H6 from unit MB4 yield YCM2o ages of 272 ± 1 Ma relatively broad peak at 339-264 Ma, which gives a YCM2o age of and 269 ± 2 Ma, respectively. The YCM2o age of tuffaceous sandstone 279 ± 2 Ma. sample M13 from unit MB5 is 263 ± 2 Ma. Sandstone sample H9 from K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 0.4 （A） B 0.3 1.2 wt % OIA (wt. OLA 0.2 0.8 ACM CIA TiO2 0.1 PCM_ 0.4 PCM' ACM- 10 10 Fe2Os+MgO (wt.%) Fe2O3+MgO (wt.%) 200 10 (C) (D) 150 ( 100 50 decreasing volcanic input OB 0.1 10 15 20 10 100 Sc (ppm) Zr/Sc La (E) sandstone ACM Moribu Formation MB4 MB5 Oguradani Formation O G MB2 MB3 Mizuyagadani Formation PCM CI andesite OIA Sorayama Formation Th Sc Fe MV 80 Mg 6.Discussion fragments of different origins (Ehiro et al., 2016; Tsukada, 2003; Tsu- kada et al., 1999). Our approach involves the precise identification and 6.1. Reconstruction of the Permian volcano-sedimentary sequence of the correlation of common lithostratigraphic features among the Paleozoic HidaGaienbelt AI petrograpic, geochemical, and detrital zircon sedimentary provenance. The Upper Paleozoic stratigraphy of the Hida Gaien belt has been The present study proposes the following reconstructed Permian stra- established in some area, including Moribu, Fukuji, and Ohno, as tigraphy (Fig. 14A), which provides a basis for better interpretation of described in Section 2. However, stratigraphic correlation and inter- the sedimentary history of this belt. The geologic time scale follows that pretation of the sedimentary history throughout this belt remain of Gradstein et al. (2020). controversial (e.g., Ehiro et al., 2016; Tazawa, 2001a,b; TAZAWA, 2002; The oldest rocks of the Permian volcano-sedimentary sequence are Tsukada, 2003) because age-diagnostic fossils have been reported from characterized by felsic tuff and minor intercalations of clastic rocks of very few horizons. In addition, differences in local lithofacies and Early Permian age, based on radiolarian assemblages and zircon U-Pb shallow-marine faunas have not been considered in terms of the lateral ages (Fig. 14A). These lithologies occur in unit MB1 of the Moribu variation and the controls of rock type on the abundance of fossil- Formation and in the lower part of the Mizuyagadani Formation. Ac- bearing horizons. As such, it has been suggested that the Paleozoic cording to Kurihara and Kametaka (2008), the lower part of the Miz- evolution of the Hida Gaien belt involved the assembly of tectonic uyagadani Formation can be correlated with the P. u-forma m. I 15 K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 Moribu Formation zoning in CL images (Fig. 12). Therefore, it is likely that volcanic activity was coeval with deposition of the radiolarian-bearing sediments in these MB1 R1 (No. 37) MB3 M5 (No. 1) 269 ± 6 Ma strata. 302 ± 12 Ma The overlying upper Lower to Midle Permian strata are units MB2 and MB3 of the Moribu Formation and the middle to upper parts of the pleochroism ranging from yellow to gold that Kubergandian to Midian (Kungurian to Capitanian; Ueno, 2006), has been recognized in units MB2 and MB3 of the Moribu Formation (Tazawa et al., 1993; Niwa et al., 2004). Although no age-diagnostic fossils are present in the middle to upper parts of the Mizuyagadani Formation, zircon grains from sandstone sample MZ5 yield a YCM2o age R6 (No. 84) 50 μm 50 μm of 279 ± 2 Ma (Fig. 13B), which corresponds to the late Early Permian 260 ± 7 Ma (Kungurian). Units MB2 and MB3 of the Moribu Formation and the MB4 M12 (No. 11) M13 (No. 2) middle to upper parts of the Mizuyagadani Formation are similar in 260 ± 8 Ma 256 ± 6 Ma terms of lithostratigraphy in that their lower sections are dominated by conglomerates, which are overlain by upper sections comprising alter- nating sandstones and mudstones (Figs. 4 and 5A). Thus, the deposi- tional ages of these strata are similar, and the strata share common lithologic features. The Mizuyagadani Formation is conformably overlain by the Sor- ayama Formation (Figs. 3A and 5A) which contains Middle Permian fusulinids (Tsukada et al., 1999). The Sorayama Formation is composed 50 μm 50 μm Epidote is common in Zone B, but sporadic in and MB3 of the Moribu Formation, which are coeval with the Sorayama MB5 H9 (No. 108) H9 (No. 98) Formation, contain volcaniclastic sandstone with abundant angular to 481 ± 13 Ma 426± 10 Ma subangular fragments of andesite (Figs. 6E and 7A). Considering their similar compositions and ages, the andesite of the Sorayama Formation is a potential source of volcaniclastic material for the Middle Permian strata, such as units MB2 and MB3 of the Moribu Formation and the Oguradani Formation, suggesting that the volcanic and clastic rock se- quences were originally intercalated within the sedimentary basin in which the above formations were deposited (Fig. 14A). The sandstone samples from units MB3 and MB4 of the Moribu H9 (No. 200) 50 μm Formation (M5, R6, and M12) yield YCM2o ages of 262±2 Ma, 440 ± 10 Ma 270 ±1 Ma, and 272±1 Ma, respectively (Fig. 13A). These detrital zircon U-Pb ages indicate early to late Middle Permian (Road- ian-Capitanian) deposition. Two YCM2o ages of 264 ± 1 Ma (sample H1 Mizuyagadani Formation of unit MB3) and 269 ± 2 Ma (sample H6 of unit MB4), which were lower part MZ1 (No. 21) upper part MZ5 (No. 22) recalculated from the data of Suzuki et al. (2019) in the northern part of 278± 7 Ma 276 ± 7 Ma the Moribu area, are similar to those in the southern part of the area (Figs. 2B and 13A). Unit MB3 contains Middle Permian brachiopod fauna consisting of Boreal/bi-temperate/Tethyan species (Tazawa, 2001b). A similar mixed fauna is also recognized in the upper part of the Oguradani Formation (Tazawa and Matsumoto, 1998), suggesting that the formation can be correlated with unit MB3. Zircons from unit MB5 of the Moribu Formation (M13) yield a late Middle Permian (Capitanian) age of 263 ± 2 Ma (YCM2o). In the northern part of the Moribu area, unit 50 μm 50 μm MB5 yields a YCM2o age of 256 ±1 Ma, indicating the early Late Permian (Wuchiapingian) (Fig. 13A). As described in Sections 3.1. and 3.2, these Capitanian-Wuchiapingian strata of the Hida Gaien belt are Fig. 12. Cathodoluminescence images of selected zircon grains from felsic tuffs characterized by clastic rocks, such as sandstone, mudstone, and alter- and sandstones of the Moribu and Mizuyagadani formations. Circles are 30 μm nating sandstone and mudstone. in diameter and show analysis spots. Previous studies have proposed that the Permian successions in the Moribu and Fukuji areas have different origins (Ehiro et al., 2016; ly to Al, and Mg tends to decrease with increasing Tsukada, 2003; Tsukada et al., 1999). The reason is that the former and the A. sinuata Range Zone (Kungurian) of Ishiga (1986). Unit MB1 of yields Monodiexodina and a brachiopod fauna which has a close affinity the Moribu Formation yields very poorly preserved radiolarians (Umeda to that from Inner Mongolia, whereas the latter yields Middle Permian and Ezaki, 1997), from which ages cannot be reliably determined. fusulinids that have been reported from South China (Tsukada, 2003; YCM2o ages from felsic tuff samples from unit MB1 of the Moribu For- Tsukada et al., 1999). In general, such a comparison requires a careful mation (R1) and the lower part of the Mizuyagadani Formation (MZ1) evaluation of the contemporaneity between the compared strata, the pict several features. The Si value in the Z-site is ecology of fossil species, and the lithological dependence of fossil the late Early Permian (Kungurian). These isotopic ages are consistent 1977; Liou,1979; Evarts & Schiffman, 1983). The with the radiolarian biostratigraphy. The felsic tuffs in unit MB1 of the the faunas of the two areas. In particular, the discussion by Tsukada Moribu Formation include abundant volcanic glass shards (Fig. 6A and et al. (1999) on the origin of the Permian succession in the Fukuji area is B). In addition, felsic tuff samples (R1 and MZ1) from both formations absent. Such an iron-rich composition may be are characterized by abundant magmatic zircons with clear oscillatory 16 K.Suzuki and T.Kurihara JournalofAsianEarthSciences 219(2021)104888 (A) Moribu Formation 282±1 R1 (MB1) 262± 2 M5 (MB3) (100%) N=39/55 19 (51 ± 18%) N = 49/84 ative 12 Prob Prob 279 ± 3 8 Z (49± 18%) ability 270±1 R6 (MB3) Relative 264 ± 1 : H1 (MB3) Relati (68.4 ± 6.9%) N= 71/91 32 (92.6 ± 4.4%): N= 77/107 wnN Proba 297 ± 3 16 289± 6 (31.6± 6.9%) (7.4± 4.4%) (1851 Ma) 272±1 M12 (MB4) Rel 269±2 H6 (MB4) Relative (80.7 ± 6.3%) N=50/69 19 (53 ± 15%) N=65/116 ? ber 14 14 Pro 299 ± 3 obability 320±6 (19.3 ± 6.3%) (43±15%)(4±21%) 263±2 M13 (MB5) 256±1 H9 (MB5) (38.5±8.9%) N = 46/60 28 (41.3 ± 7.4%) N = 120/224 iber 73±2 21 3 ± 7.3%) Pro 284 ± 2 Z excluding one Precambrian age (61.5 ± 8.9%) 01±3 (2353 Ma) ±10%) :378 426 481 562 360400440480 520 560 360400 440480 520 200 240 280 320 600 200 240 280 320 560 600 206Pb/238U age (Ma) 206Pb/238U age (Ma) (B) Mizuyagadani Formation 283±2 MZ1 (lower part) MZ5 (upper part) 279±2 (100%) N = 19/27 20 (42.9 ± 9.8%) N=68/80 oer 15 Prob 10 302±3 (57.1 ± 9.8%) 403 200240280 320 360 400 440480 520 560 600 200 240 280 320 360 400 440 480 520560 600 206Pb/238U age (Ma) 206Pb/238U age (Ma) youngest cluster age of cluster age components of Mixture Models (YCM2o) Mixture Models uyagadani formations. The data of samples H1 and H6 are used from those of samples 0610-05 and 0529-63 reported by Suzuki et al. (2019), respectively. The age distribution for sample H9 reflects a combination of data from this study and from sample 0818-06 of Suzuki et al. (2019). There is little basis for inferring the origin of the Moribu and Fukuji areas 6.2.Provenance of the Permian volcano-sedimentary sequence in the based on differences in their fossil faunas. In contrast, lithostratigraphic Hida Gaienbelt and age data of the present study for the Moribu, Mizuyagadani, Sor- ayama, and Oguradani formations show close correlation, and the for- Provenance analyses have been conducted for samples of Permian mations share common lithologic features. Therefore, it is reasonable to shallow-marine sedimentary rocks from the Hida Gaien, South Kitakami, suggest that the Permian formations of the Hida Gaien belt were origi- and Kurosegawa belts (e.g., Yoshida et al, 1994; Yoshida and Tazawa, nally deposited in a single sedimentary basin. As such, the reconstructed 2000; Yoshida and Machiyama, 2004; Hara et al., 2018a). Of these Permian volcano-sedimentary sequence of the present study was domi- studies, Yoshida et al. (1994) and Yoshida and Machiyama (2004) nated by volcanic and volcaniclastic rocks and sandstones rich in vol- showed a clear change in provenance for the Permian strata in the South Kitakami belt from an undissected arc source area during the Early to gradually to clastic rocks during the late Middle to Late Permian mid-Middle Permian to an uplifted arc basement source area during the (Fig. 14A). late Middle to Late Permian. Trench-fill sediments of Middle-Late e s show similar trends in terms of input from volcanic/plutonic rock detritus, as recognized from geochemical analyses and modal 7 K.Suzuki and T.Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 Suolun- (A) Hida Gaien belt (B) Jilin (C) Xi Ujimqin Changhsingian Yangjiagou Linxi Wuchia 工 MB5 Formation Formation D pingian 6H (Moribu Fm.) Capita- nian MB4 e (Moribu Fm.) Fanjiatun Zhesi Middl M5 H1 M13 Formation Formation Wordian 0.00 Permian M 工 Roadian R6 H6 M12 Miz Daheshen Kungu- 0 Formation Dashizhai rian Formation Fm MZ5 ani R1 么 Artins MZ1 Shoushangou Q kian Formation M Sakma- dated samples rian (see Fig. 13) Shizuizi Amushan lowermost part of Formation Asselian Formation the Mizuyagadani Fm mudstone, alternating volcaniclastic sandstone nassivesandstone and mudstone volcanicrocks sandstone and mudstone felsic tuff, tuffaceous nglomerate sandstone and mudstone calcareousclasticrocks limestone youngest cluster age Boreal/bi-temperate/ adiolarians with 2o uncertainty (YCM2o) Tethyan brachiopods Monodiexodina Fig. 14. Stratigraphic reconstruction and zircon U-Pb ages for Permian strata of the (A) Hida Gaien belt, compared with coeval strata of the (B) Jilin and (C) Suolun-Xi Ujimqin areas. The geological time-scale is based on Gradstein et al. (2020). Stratigraphic information in (B) and (C) is from Shen et al. (2006) and Han et al. (2017), respectively. compositions (Takeuchi et al., 2008; Hara et al., 2018a; Zhang et al., period has also been recognized in the Nagato tectonic belt and the 2018; Ohkawa et al., 2021). Kurosegawa belt of Southwest Japan (Hada et al., 1992; Kametaka, For the Hida Gaien belt, Yoshida and Tazawa (2000) conducted 2006; Kurihara and Kametaka, 2008; Kuwahara et al., 2009). In the provenance analysis in the type section of the Moribu Formation, which eastern part of East Asia during the late Early Permian (ca. 280 Ma), partly corresponds to units MB2-MB4 in the present study, based on active arc volcanism was initiated in the southeastern margin of the petrographic signatures and major-element-oxide contents. Those au- South China block (the Cathaysia block) by subduction of the Paleo- thors recognized that abundant fragments of andesitic volcanic rocks are Pacific oceanic plate (Li et al., 2006; Li et al 2012). Coeval magma- present in the lower part of the formation and that the abundances of tism is also recognized in the Jiamusi-Khanka block (Wu et al., 2011; fragments of various plutonic and volcanic rocks increase upward Yang et al., 2015). The Maizuru belt (Yakuno ophiolite) of Southwest through the stratigraphy. In the present study, we integrated sandstone Japan also occurred in an oceanic island arc-back arc system in the petrographic and geochemical data with zircon U-Pb ages to reconstruct western Paleo-Pacific Ocean (e.g., Hayasaka et al., 1996; Ishiwatari, the entire Permian volcano-sedimentary sequence of the belt (Fig. 14A). 1990; Ichiyama and Ishiwatari, 2004; Kojima et al., 2016). This means Here, using these data, we discuss the change in provenance that that during the Early Permian, arc volcanism related to oceanic plate occurred during the Permian (Fig. 16A). subduction was widespread across the eastern margin of East Asia and The Lower Permian strata of the Hida Gaien belt (unit MB1 of the that the provenance of the Lower Permian sedimentary strata of the Hida Moribu Formation and the lower part of the Mizuyagadani Formation) Gaien belt was strongly influenced by felsic magmatism of the arc. consist mainly of felsic tuffaceous rocks. This suggests that felsic arc The upper Lower to Middle Permian strata (unit MB2 of the Moribu volcanism occurred during the Early Permian (Figs. 14A and 16A). Although the Lower Permian of Japan is typified by shelf-carbonate are characterized by abundant andesitic volcanic rock fragments facies (i.e., the Sakamotozawa Formation of the South Kitakami belt; (Figs. 6E and 7A, F), as confirmed by major-element-oxide contents e.g., Kanmera and Mikami, 1965a; Kanmera and Mikami, 1965b), the (Table 1). In addition, the REE patterns for these strata show depletion in widespread occurrence of similar tuffaceous lithologies during this time LREEs relative to UCC, with no negative Eu anomalies (Fig. 10A), 18 K.Suzuki and T.Kurihara JournalofAsianEarthSciences219(2021)104888 450(A) Permianarcvolcanism Hida Gaien belt Relative Probability (280-250 Ma) Number 350- N= 546 This study 250- 500-400Ma & Suzuki et al. (2019) 150 graniticbasement 1851Ma 2353Ma 50 100 Jilin area Relative Probability 80 N = 176 hber 60- Wang et al. (2015) 40 20- (C) 250- Inner Mongolia Relative Probability 200 N = 731 lumbe 150 Chen et al. (2016) & Han et al. (2017) 100- 50- (D) Southeastern margin of Relative Probability 25 South China block the South China block Number 20 N = 183 15- Hu et al. (2014) 10- 60 Southwestern margin of Relative Probability North China block 50 the North China block Number 40- N = 421 30- Liang et al. (2020) 20- 10- 250 500 750 0 206Pb/238U age (Ma) Fig. 15. Comparison of detrital zircon U-Pb age spectra of the Permian strata for different regions of East Asia. (A) Hida Gaien belt (Suzuki et al., 2019 and this study); (B) Jilin area (Wang et al., 2015); (C) Inner Mongolia (Chen et al., 2016; Han et al., 2017); (D) southeastern margin of the South China block (Hu et al., 2014); (E) southwestern margin of the North China block (Liang et al., 2020). K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 felsic to intermediate (A) volcanic activity volcanism granitic detritus basement uplift Permian arc basement pre-Permian strata source rock variability Ordovician-Devonian arc basement high 280 270 260 (Ma) Early Middle Late Permian (B) Early-Middle Permian Songliao-Xilinhot block Suolun-Xi Ujimqin Jilin Jiamusi-Khanka-Bureya block Ocean Paleo-Pacific Ocean North China HG SK KS Paleo-Tethys Ocean South China AK MZ 500-400 Ma granites and cratons arc volcanic activity trench-fill sediments microcontinents Fig. 16. (A) Schematic model of the sedimentary supply system and magmatic arc development of the Hida Gaien belt during the Permian, and (B) reconstruction of the tectonic setting of Northeast Asia for the Early-Middle Permian. HG: Hida Gaien belt; SK: South Kitakami belt; KS: Kurosegawa belt; MZ: Maizuru belt; AK: suggesting that the detritus was derived from basaltic to andesitic vol- evolve since the Early Permian (Fig. 16A). canic rocks (e.g., McLennan, 1989). Diagrams of Al2O3/ Sandstones of the Middle Permian strata (unit MB3 of the Moribu SiO2-(Fe2O3 + MgO), TiO2- (Fe2O3 + MgO), and La-Th-Sc indicate that Formation) and the Oguradani Formation contain not only volcanic the tectonic setting of the source area was most likely an oceanic island detritus but also more quartz and feldspar grains than stratigraphically arc (Fig. 1lA, B, and E). With respect to the determination of the main lower strata, such as unit MB2 (Figs. 7B, C, and 8). The SiO2 contents of source rock, data from these strata plot around the Paleozoic andesite in samples from unit MB3 are relatively high (Table 1), and the REE pat- Th/Sc-Zr/Sc and La-Th-Sc diagrams (Fig. 11D and E). As mentioned in terns are largely similar to that of UCC, although the Eu values and Section 6.1, the potential source of the andesitic detritus is the Middle HREEs are slightly variable (Fig. 10). In Th/Sc-Zr/Sc and La-Th-Sc Permian Sorayama Formation, which consists mainly of andesitic vol- diagrams, data for this unit plot between the Paleozoic andesite and canic and volcaniclastic rocks (Fig. 14A). Middle Permian sandstones Paleozoic felsic volcanic rocks fields (Fig. 11D and E). This indicates a rich in andesitic volcanic detritus have also been recognized in other provenance transition from an immature volcanic arc to a mature arc terranes, such as the South Kitakami belt (Yoshida and Machiyama, during the Middle Permian. 2004); however, andesitic rocks have not been reported for this period. Sandstones of the The present study has proposed a possible source of the andesitic itanian-Wuchiapingian) strata (units MB4 and MB5 of the Moribu For- detritus, and in the Hida Gaien belt, the andesites of the Sorayama mation) are more feldspathic than those of units MB2 and MB3 (Figs. 7D, Formation are a very likely source. This andesitic volcanic source played E and 8). The SiO2 contents of the sandstones in these units are higher an important role in the sediment supply for Middle Permian deposition, than stratigraphically lower strata (Table 1), and their REE patterns are with the andesites being derived from an arc that had continued to characterized by negative Eu anomalies similar to the UCC value 20 K. Suzuki and T. Kurihara Journalof Asian Earth Sciences 219(2021）104888 (Fig. 10A). On Al2O3/SiO2- and TiO2- (Fe2O3 + MgO) and La-Th-Sc 2015). The present study uses the geotectonic units of Chen et al. (2017), diagrams, data from these units plot in and around the continental island n s n s arc field (Fig. 11A, B, and E). Contents of V and Sc, which are influenced orogenic belt (SOB); (2) the Solonker Suture Zone / Hunshandake by the input of volcanic detritus in a sedimentary basin, decrease Block (SSZ/HB); (3) the northern early to mid-Paleozoic orogenic belt gradually upward through the stratigraphy (Fig. 11C). On the basis of (NOB); (4) the Erenhot-Hegenshan ophiolite belt (EHOB); and (5) the Th/Sc-Zr/Sc and La-Th-Sc diagrams, the main rocks exposed in the Uliastai Continental Margin (UCM; Fig. 1). In the tectonic subdivision by source area are inferred to have changed gradually from andesite to Eizenhofer and Zhao (2018), the SOB, SSZ/HB, NOB, and EHOB corre- granite (Fig. 11 D and E). spond roughly to the Bainaimiao arc belt, the Solonker Suture, the The U-Pb ages of detrital zircons provide further information on this Baolidao arc belt, and Hegenshan back-arc belt, respectively. The eastern extension of the SOB, SSZ/HB, and NOB is truncated by the of the present study (sample H9) yields detrital zircons of Ordovician to Songliao Basin, in which Mesozoic terrestrial cover is widely distributed. early Late Devonian age (480-378 Ma; Fig. 13A). The Permian volcanic In the area east of the Songliao Basin (Northeast China), the eastern arc of the Hida Gaien belt must have been generated upon an Ear- extension of the Bainaimiao arc belt, the Solonker Suture, and the ly-Middle Paleozoic arc basement, given that the Upper Paleozoic strata are underlain by Ordovician to Devonian formations that were deposited responding to the SOB, SSZ/HB, and NOB, respectively. Furthermore, in an arc-related basin that was strongly influenced by arc volcanism. Liu et al. (2017) proposed the Songliao-Xilinhot block represents a The occurrence of these Ordovician-Devonian zircons indicates that the granitic rock basement that corresponds roughly to the SSZ/HB and 500-400 Ma arc basement had been uplifted and exposed during the NOB. It is considered that this block consists mainly of Early Cambrian Late Permian (Fig. 16A). Similar occurrences of older detrital zircons to Silurian-Devonian granitic rocks (e.g., Chen et al., 2000; Jian et al., within Upper Permian strata are widely recognized in the South Kita- 2008). The representative Permian formations of the eastern CAOB are Formation of the South Kitakami belt yields detrital zircons of 530-497 distributed mainly in the Jilin area, east of the Songliao Basin (Northeast and 468-445 Ma in age (Okawa et al., 2013), and sandstones within the China), and in several areas of the SSZ/HB, NOB, and EHOB west of the Upper Permian accretionary complex of the Kurosegawa and Akiyoshi Songliao Basin (Inner Mongolia; Fig. 1). Applying the tectonic division belts contain 500-400 Ma detrital zircons (Hara et al., 2018a; fig. 8 in of Liu et al. (2017), the distribution zone of the Permian rocks of the Jilin Zhang et al., 2018; Ohkawa et al., 2021). area corresponds approximately to the SSZ/HB and NOB. According to In summary, the provenance of the Permian volcano-sedimentary Shen et al. (2006) and Han et al. (2017), the Lower to Middle Permian sequence of the Hida Gaien belt changed gradually from a source area formations in the Jilin area and in the Suolun and Xi Ujimqin areas comprising an immature undissected arc with active felsic to interme- (EHOB) are characterized by abundant volcanogenic components. In diate volcanism during the Early to Middle Permian to a mature arc addition, these strata contain abundant Permian brachiopods and during the late Middle-Late Permian. The uplift of the matured arc and fusulinids. diminishing volcanic activity during the Late Permian provided sedi- Permian rocks in the Jilin area are subdivided into the Shizuizi, ment from the root of the matured Permian arc and the Ordovi- Shoushangou, Daheshen, Fanjiatun, and Yangjiagou formations in cian-Devonian arc basement. This occurred widely across the magmatic ascending stratigraphic order (Shen et al., 2006; Fig. 14B). Among these arc source areas of the shallow marine and trench-fill sedimentary rocks formations, the Daheshen Formation is composed of abundant volcanic of the major Paleozoic terranes of Southwest and Northeast Japan. and pyroclastic rocks such as andesite, rhyolite, and felsic to interme- diate tuff, with a few limestone intercalations. On the basis of the 6.3.Comparison between the Permian strata of the Hida Gaien belt and occurrence of fusulinids, including Monodiexodina, the Daheshen For- the eastern CAOB mation is inferred to be Kubergandian in age (Guo et al., 1992; Ueno and Tazawa, 2003). The overlying Fanjiatun Formation consists of The provenance of the Permian strata of the Hida Gaien belt and mudstone, volcanic rocks, and volcaniclastic sandstone interbedded other Paleozoic terranes of Japan was strongly influenced by the evo- with limestone. The brachiopod fauna in these rocks (Guo et al., 1992) lution of a magmatic arc. For further evaluation of the spatial scale of suggests that the Fanjiatun Formation is early Middle Permian (Road- this magmatic-arc provenance, paleobiogeographic correlation between ian-early Wordian) in age. Permian faunas in Japan and Northeast China-Inner Mongolia (eastern In the Suolun and Xi Ujimqin areas, Permian strata are subdivided CAOB) is essential. A Middle Permian brachiopod fauna from the Hida into the Amushan, Dashizhai, Zhesi (or Jisu Honguer), and Linxi for- Gaien and South Kitakami belts, comprising a combination of Boreal/bi- mations in ascending stratigraphic order (Han et al., 2017). The temperate/Tethyan species, has been recognized broadly across Inner Dashizhai and Zhesi formations are equivalent to the Daheshen and Mongolia, Northeast China, and South Primorye (Far East Russia) (e.g., Fanjiatun formations in the Jilin area, respectively. The Dashizhai For- Tazawa, 1998; 2001a; 2002; Shi, 2006). In addition, Monodiexodina—a mation consists of lava and volcaniclastic rocks containing rhyolite and Permian antitropical fusulinoidean genus that occurred in midle andesite, which are slightly metamorphosed. This formation is Roadian paleolatitude areas in the northern and southern hemispheres—has been to Wordian in age based on brachiopod assemblages including Yakov- reported from the same regions (Ueno, 2006). Therefore, comparison of levia (Jin et al., 2000). The overlying Zhesi Formation is composed of the Permian lithostratigraphy and provenance of the Hida Gaien belt mudstone, volcaniclastic sandstone, and limestone. On the basis of with strata in Northeast China-Inner Mongolia, with the two areas being brachiopods such as Yakovlevia and Spiriferella, this formation is inter- connected prior to the opening of the Sea of Japan, should provide a preted to be Wordian in age (Manankov, 1999). better understanding of their regional geologic evolution and tectonic Detrital zircon U-Pb analyses have been conducted for the Fanjiatun relationship with proto-Japan and the eastern CAOB. In the following and Zhesi formations (Wang et al., 2015; Chen et al., 2016; 2017; Han paragraphs, we provide an overview of the stratigraphy and provenance et al., 2017). In the Jilin area, Wang et al. (2015) determined that of the Permian strata of Northeast China-Inner Mongolia, and discuss feldspathic sandstones from the Fanjiatun Formation have a dominant their similarities to and differences from the strata of the Hida Gaien belt age peak at 264 Ma (Capitanian), together with a number of grains with in terms of lithostratigraphy, provenance, and detrital zircon ages Early to Middle Paleozoic ages of 529-486 and 441-310 Ma; however, (Figs. 14 and 15). Proterozoic zircons are rare (Fig. 15). A similar age peak at 265-262 Ma The geotectonic framework of the eastern CAOB has been well has been reported from the underlying and overlying Permian strata (the studied in the area west of the Songliao Basin in Inner Mongolia (e.g., Shoushangou and Yangjiagou formations). Wang et al. (2015) inter- Miao et al., 2008; Jian et al., 2010; Heumann et al., 2012; Zhang et al., preted that the Permian zircons were possibly derived from a source 21 K. Suzuki and T. Kurihara Journalof Asian Earth Sciences 219(2021）104888 related to continent-continent collision between the continental margin arc. According to Eizenhofer and Zhao (2018), the Jilin area belonged to of the North China block and a supposed microcontinent. For the Ear- the East Asian Pre-Pacific Province, which was strongly influenced by ly-Middle Paleozoic zircons, Wang et al. (2015) suggested that the oceanic plate subduction, and thus this interpretation is reasonable and source area was a magmatic arc of Ordovician-Early Devonian age (i.e., n u the Zhangjiatun island arc), rather than the North China block. In the In addition, on the basis of the ages of granitic rocks in the Jiamusi-- comprehensive discussion of the provenance of the eastern CAOB by Khanka block (ca. 523-515 Ma; e.g., Wilde et al., 2003; Yang et al., Eizenhofer and Zhao (2018), the “East Asian Pre-Pacific Province" was 2015), this block is considered to be a possible source of Early Paleozoic proposed for the region east of the Songliao Basin, including the Jilin zircons of the Hida Gaien belt and the Jilin area. area and the neighboring Jiamusi-Khanka-Bureya block. According to The Zhesi Formation, which is distributed further to the west than Eizenhofer and Zhao (2018), the Permian sedimentary rocks are char- the Songliao Basin, yields abundant older grains, including 500-400 Ma acterized by abundant Early-Middle Paleozoic detrital zircons and rare zircons, with Permian zircons (Chen et al., 2016; 2017; Han et al., 2017; Proterozoic zircons. Fig. 15). As reviewed by Liu et al. (2017), 500-400 Ma granitic rocks are Chen et al. (2016; 2017) reported detrital zircon U-Pb ages from common in the western part of the SSZ/HB and NOB as the basement Carboniferous-Permian strata of the SSZ/HB, NOB, and EHOB. In the represents part of the Songliao-Xilinhot block. This suggests that the western part of the SSZ/HB, lithic sandstone (sample NM07-46) from the degree of uplift of the older granitic rock basements varied in the source Zhesi Formation in the Sonid Right Banner (Sonid Youqi) area is char- areas for the Zhesi Formation and for the Permian strata of the Hida acterized by a dominant age peak at 416 Ma, subsequent peaks Gaien belt and the Jilin area, although the fauna in the Zhesi Formation (303-278 Ma) close to the depositional age (fig. 7 in Chen et al., 2016), suggests the same mid-latitude paleobiogeographic province as those of and minor Proterozoic zircons. A similar age distribution is recognized the Hida Gaien belt and the Jilin area. It is therefore likely that the Zhesi in the Zhesi Obo area (sample NM08-85; fig. 8 in Chen et al., 2016). oo According to the results of Chen et al. (2016; 2017), the sandstones of basement in terms of sediment provenance compared with the Permian the Zhesi Formation are rich in Early-Middle Paleozoic and Permian strata in the Hida Gaien belt and the Jilin area (Fig. 15). zircons, and their abundances differ depending on the region and hori- zon. For example, sandstone (sample NM12-178) in the lower part of the 6.4. Implications for the tectonic setting and paleogeography of the Hida Zhesi Formation in the Sonid Right Banner area is dominated by 445 Ma Gaien belt and the eastern CAOB zircons with rare Permian zircons (fig. 10b in Chen et al., 2017), whereas sandstone (sample NM12-183) in the upper part of the formation has a The Permian tectonic setting and paleogeography of proto-Japan and dominant age peak at 280 Ma rather than 445 Ma (fig. 10d in Chen et al., the eastern CAOB in Northeast Asia have traditionally been understood 2017). In the Sonid Left Banner (Sonid Zuoqi) area, sandstone (sample as occurring within different frameworks. For example, proto-Japan is tween actinolite and hornblende as shifting toward 361 Ma grains, and sandstone (sample NM10-29) of the EHOB has broad Paleo-Pacific Ocean (e.g., Maruyama et al., 1997; Isozaki et al., 2010). In age peaks at 483-376 and 338-275 Ma. Han et al. (2017) reported contrast, the eastern CAOB is regarded as an assembly of various similar zircon age spectra from the EHOB for the Suolun and Xi Ujimqin geologic units that were related to multiple collisions/amalgamations of areas, with dominant age peaks at 515-511 and 323-272 Ma. Taken microcontinents and magmatic arcs between the North China and Si- together, sandstones of the Zhesi Formation are characterized by an In the medium-grade part of Zone C, both abundance of older grains, such as Early-Middle Paleozoic zircons, with zircon ages mentioned above, it is concluded that Permian magmatic arc Permian zircons (Fig. 15). In addition, considering that the Zhesi For- and Early-Middle Paleozoic arc basements were widely distributed, mation is Wordian (268-265 Ma) in age, as inferred from brachiopods, given that they represent the source(s) from which the sediment even the Permian peaks are not syn-depositional, but rather slightly deposited in basins of the Hida Gaien belt to the eastern CAOB was older than the depositional age as estimated by the biostratigraphy. derived during the Permian. As noted by Wang et al. (2015), continental These data are consistent with a mixed provenance of older detritus and blocks of Precambrian age, such as the North China block (e.g., Liang a felsic source inferred from the sandstone geochemistry of Han et al. et al., 2020), which lies close to Japan and the eastern CAOB in the (2017). present day, made little contribution to the sediment supply (Fig. 15). On the basis of the biostratigraphic and isotopic ages discussed Isozaki et al. (2014; 2015; 2017) argued that 500-400 Ma granitic above, the Permian volcano-sedimentary sequence of the Hida Gaien rocks, which were the main constituents of the Early-Middle Paleozoic belt can be correlated with those of the Jilin, Suolun and Xi Ujimqin arc basements, are distributed continuously from the eastern continental areas (e.g., Shen et al., 2006; Han et al., 2017; Fig. 14). In terms of margin of the South China block through Southwest and Northeast grade of metamorphism, ranging from 0.10 (No. Japan to the Jiamusi-Khanka-Bureya block in Far East Asia. They pro- change from volcaniclastic/pyroclastic to clastic rocks. Felsic tuff and n r e i ( ns rr, a psod andesite are common in the lower part of the Mizuyagadani Formation, margin along a single subduction zone in the Paleo-Pacific Ocean might the Sorayama Formation, and the Daheshen and Dashizhai formations. have occurred from the South China block to the eastern end of the This suggests that felsic to intermediate arc volcanism during the Early CAOB. However, 500-400 Ma granitic rocks are widely distributed in to early Middle Permian contributed to the development of the Hida Central and East Asia(e.g., Rojas-Agramonte et al., 2011; Lin et al., 2013 Gaien belt and the eastern CAOB. In terms of sandstone provenance, the and references therein), and it is unlikely that all of the granitic rocks of uppermost part of the Moribu Formation (sample H9 from unit MB5) that age were originated entirely from the South China block. In addi- yields a dominant Late Permian age peak, several Early-Middle Paleo- tion, the Permian strata of the Hida Gaien belt and Northeast China- zoic zircons with ages of 480-378 Ma, and rare Proterozoic zircons -Inner Mongolia do not have clear age peaks of Precambrian zircons that (2363 and 562 Ma; Fig. 13A). This is similar to the age spectra for the suggest a strong relationship with the South China block (e.g., Hu et al. Fanjiatun Formation in the Jilin area (fig. 8b in Wang et al., 2015; 2014; Fig. 15). Taking these factors into consideration, we propose a Fig. 15). As mentioned above, Wang et al. (2015) considered that the new reconstruction of the tectonic setting, as shown in Fig. 16B. The N-S Permian strata in the Jilin area consist of sediments derived from con- arrangement of the Hida Gaien, South Kitakami, and Kurosegawa belts tinent-continent collision. However, the similarity in lithostratigraphy follows that proposed by Tazawa (200la). The tectonic setting of sub- crest about Al'V=0.9. It seems that, with increas- - pe soa pus e sd oo p sz un et al. (1983) regarded the compositional gap be- zuru belts, characterized by two juxtaposed arc-trench systems, are tectonic seting of the Jilin area; namely, the syn-depositional zircons referred to Hayasaka et al. (1996). The location of the subduction zone compositional gap can be defined. With increasing in which the accretionary complex of the Akiyoshi Belt was formed is 22 K. Suzuki and T. Kurihara Journal of Asian Earth Sciences 219 (2021) 104888 based on Zhang et al. (2018). In the model, we assume that a Permian related basins along a single subduction zone is inferred to have magmatic arc acting as the sediment source for the Permian strata of the extended to at least the Jilin area in the eastern CAOB. Hida Gaien belt and the eastern CAOB formed on 500400 Ma arc basements along a single subduction zone in the western Paleo-Pacific CRediT authorship contribution statement Ocean. With regard to the Permian of the Jilin area, based on the lith- ostratigraphic and age similarities to the Hida Gaien belt mentioned Keisuke Suzuki: Conceptualization, Methodology, Software, Formal above, the volcanic rocks and syn-depositional zircons would have been analysis, Investigation, Writing original draft, Writing review & derived from the magmatic arc along the 500400 Ma arc basement. It is editing, Visualization. Toshiyuki Kurihara: Investigation, Supervision, highly possible that the Permian strata of the Hida Gaien belt and the Project administration, Writing review & editing. Jilin area accumulated in the same tectonic setting along the single subduction zone.Furthermore, the oceanic plate subduction is likely to have occurred during the Early Permian, according to evidence of arc Declaration of Competing Interest magmatism from the South China, Japan, and JiamusiKhanka margins (e.g., Li et al., 2012; Kojima et al., 2016; Ishiwtari et al., 2016; Yang The authors declare that they have no known competing financial et al., 2015). In response to the active subduction and subsequent interests or personal relationships that could have appeared to influence compression, the EarlyMiddle Paleozoic arc basements were gradually the work reported in this paper. uplifted and acted as the source of the 500400 Ma detrital zircons. The SongliaoXilinhot block in the eastern CAOB, corresponding to Acknowledgments the SSZ/HB and NOB of Chen et al. (2017), is composed of 500400 Ma granitic rocks (e.g., Chen et al., 2000; Jian et al., 2008; Liu et al., 2017). We thank Dr. Hayato Ueda for his help with zircon U-Pb dating. We Therefore, the 500400 Ma arc basement forming proto-Japan may also thank Dr. Toshiro Takahashi and Ms. Rikako Nohara-Imanaka for extend further west into the SongliaoXilinhot block. However, this their assistance with XRF and ICPMS analyses. We are grateful for the hypothesis can still not be solved only with our data, and it is unclear constructive comments from two anonymous reviewers and Dr. Paul R. how the oceanic plate subduction in the Paleo-Pacific Ocean influenced Eizenhofer, all of which have greatly improved our manuscript. This the Permian strata in the areas west of the Songliao Basin. In general, study was supported by a Fukada Grant-in-Aid 2018 from the Fukada paleogeographic reconstructions of the eastern CAOB suggest a Geological Institute, Japan and JSPS KAKENHI Grant Number complicated arrangement of subduction zones around the Solonker JP17K05690. Ocean, where the tectonic units of Inner Mongolia developed, rather than in the Paleo-Pacific Ocean (Torsvik and Cocks, 2017). The simi- Appendix A.Supplementary tables larities between the Hida Gaien belt and eastern CAOB in terms of the development of Permian volcanic rocks and the age distributions of Supplementary tables to this article can be found online at https detrital zircons suggest that the effects of oceanic plate subduction might ://doi.org/10.1016/j.jseaes.2021.104888. have extended further to the west. With these possibilities in mind, the tectonic evolution of the intersection between the orogenic belts of the References Paleo-Pacific Ocean and the CAOB should be carefully considered in future studies. Adachi, S., 1985. Smaller foraminifers of the Iehinotani Formation (Carboniferous-Permian), Hida Massif, central Japan. Sci. Rep. Inst. Geosci. Univ. Tsukuba Sec. 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vww.nature.com/scientificreport SCIENTIFIC REPORTS OPEN Groundwater helium anomaly reflects strain change during ithe 2016 Kumamoto earthquake in Received: 19 September 2016 Southwest Japan Accepted: 01 November 2016 Published: 29 November 2016 Dapeng Zhao* Geochemical monitoring of groundwater and soil gas emission pointed out precursor and/or coseismic basis. A laboratory experiment of rock fracturing and noble gas emission was conducted, but there is no quantitative connection between the laboratory results and observation in field. We report here deep groundwater helium anomalies related to the 2016 Kumamoto earthquake, which is an inland crustal earthquake with a strike-slip fault and a shallow hypocenter (10 km depth) close to highly populated areas in Southwest Japan. The observed helium isotope changes, soon after the earthquake, are quantitatively coupled with volumetric strain changes estimated from a fault model, which can be explained by experimental studies of helium degassing during compressional loading of rock samples. Groundwater helium is considered as an effective strain gauge. This suggests the first quantitative where conventional borehole strain meter is not available. Volatile element degassing from the solid Earth has been a major subject of geochemistry because it may provide mental consequence of volcanism23. Even though magmatic and hydrothermal fluxes are spectacular and have a high profile45, gas emanations related to seismic and fault activities are not negligibly small and may contribute atmospheric mass balance6-8. Geochemical monitoring of groundwater and soil gas emission has been conducted for earthquake prediction in USA, Japan, China and Italy since 1970s?. Precursor and/or coseismic anomalies of noble gases such as groundwater radon and helium/argon ratios were well documentedio. Several years prior to an earthquake in Tashkent, radon contents in deep wells increased to about double normal values. Soon after the earthquake, they returned to normal. It was the first report of radon anomaly and stimulated many following researchers. During the 1995 Kobe earthquake, continuous radon-monitoring was made at a site located about 30km northeast of the epicenter12. The radon concentration at the end of 1993 was stable at 20 Bq per liter. Radon started to increase gradually from October 1994 to the end of November 1994, reaching about 60 Bq per liter, which was three times that of the normal level. A sudden increase was seen on January 7, 1995 and then decreased on January 10, a week before the Kobe earthquake. Since then the radon content returned to the November 1993 level. It was the most significant precursor of a large earthquake. In addition, conspicuous changes in He/Ar ratios were observed at a fumarole of Mt Ontake and a mineral spring in the Byakko just before the occurrence of an inland earthquake (M 6.8) in central Japan in September 1984; the fumarole and the spring are located 9 and 50kilometers, respectively, from the epicenterl3. The anomaly was probably resulted from deep-seated flu- plausible physico-chemical basis to explain these phenomenal4. A laboratory experiment of rock fracturing and Atmosphere and Ocean Research Institute, The University ofTokyo, Kashiwa, Chiba 277-8564, Japan. 2lnstitute for Geothermal Sciences, Kyoto University, Beppu, Oita 874-0903, Japan. 3Department of Earth and Environmental Sciences, Kumamoto University, Kurokami, Kumamoto 860-8555, Japan. 4Department of Geophysics, Tohoku (email: ysano@aori.u-tokyo.ac.jp) www.nature.com/scientificreports/ 33°06'N 40'N OKKC OUKI OTMN ASO Kumamoto Futagawa Fault OTS -32°54'N Zone Main shock Kumamoto 16 April (M7.3) Foreshock 14 April (M 6.5) 32°42'N MEE AJS Hinagu Fault Foreshock 15 April (M 6.4) Zone 20km Matsushiro 32°30'N 130°18'E 130°42'E 131°06'E Figure 1. Sampling sites of deep groundwater in the Kumamoto region in Southwest Japan together with epicenters of two foreshocks (M 6.5 and M 6.4) and the main shock (M 7.3) of the 2016 Kumamoto Portal Web Site, Geospatial Information Authority of Japan (http://maps.gsi.go.jp). helium and radon emission was conducted15,16, but there is no quantitative relationship between the laboratory results and field observation. It is necessary to connect gas geochemistry, model of strain change and rock frac- turing experiment in a comprehensive study. The 2016 Kumamoto earthquake (M 7.3 in the Japan Meteorological Agency scale)17 occurred in Kumamoto City on Kyushu Island, Southwest Japan on April 2016, causing 50 fatalities, over 1800 injured and serious damage to local infrastructures. It was an inland crustal earthquake with a strike-slip fault and a shallow hypo- center (10 km depth) close to highly populated areas (see Methods). In order to study geochemical state of the Kumamoto area, we collected groundwater samples. There are noble gas data collected and measured in the same areas in August 201018. So it is possible to compare directly the helium isotopes before and soon after the Kumamoto earthquake. Based on the data, we discuss the connection between emanation of crustal helium, esti- mated volumetric strain change, and a quantitative comparison with a rock fracturing experiment. Results Helium isotopes of groundwater samples. Seven deep groundwater samples (280-1300 m) were col- lected in the Futagawa-Hinagu fault zones during April 28 to 29, 2016 (Fig. 1), eleven days after the Kumamoto main shock. Helium and neon abundances and helium isotopes were measured by conventional mass spec- trometers (Methods). The He/He ratios of samples derived from the upper mantle show a high value of 8 Ra (Ra is the *He/He ratio of 1.382 × 10-6)19, whereas those of crustal fluids are characterized by a radiogenic ratio of ~0.02 Ra20. Observed 3He/+He ratios of our samples vary from 0.623 Ra to 4.12 Ra (STable 1) and show that ground- water in the Futagawa-Hinagu fault zones contains mantle helium contributions of ~7% to ~50%. The highest ratio uss found (Fig. 1). An elevated *He/*He ratio was also observed in Mifune hot spring, close to the epicenter of a major foreshock (M 6.4) on April 15. However, the ratios are lower than the air value at distant sites from the faults, such as Tamana and Hirayama hot springs. Figure 2a shows a relationship between the approximate distances from the Futagawa-Hinagu fault to the sampling sites and the bserved He/He ratios of our samples together with those measured six years before the 2016 Kumamoto earthquake18. The overall variations are consistent before and soon after the earthquake, and there is amodest trend of decreasing ratios with distance away from the Futagawa-Hinagu fault. The same trend was first recognized in the San Andreas fault system in California2l and confirmed by later studies in theNorth Anatolian fault zone in Turkey22 and the Karaforam fault in Tibet23. Note that they are allong strike-slip faults, and a general explanation of the location-dependent variation is due to influx of mantle fluids as follows: helium with a high 3He/*He ratio exsolving from partially-melted zones in the upper mantle becomes During the uprising of the fluids, they may be diluted by radiogenic helium with a low ratio. At the distant sites from ai oo u sds ao s s s d q u desn SCIENTIFIC REPORTS|6:37939|DOl:10.1038/srep37939 www.nature.com/scientificreports/ Before earthquake (a) TS After earthquake R MFN UKI AJS KKC ： .·HRY SENW TMN -10 0 10 20 30 40 50 (Ra) 0.2 AJS (b) nge 0.0 HRY TMN.. -0.2 MFN KKC UKI -0.4 /4He -0.6 OTS e -0.8 阳 -10 0 10 20 30 40 50 (c) H P 4 OTS 3 cm3 MFN TT KKC I UKI 0 HRY Y -10 0 10 20 30 40 50 Distance from fault (km) Figure 2. Relationship between the distance from the fault to the sampling site and (a) He/*He ratios of deep groundwater before and after the M 7.3 earthquake, (b) that between the distance and the *He/*He change, and (c) that between the distance and the radiogenic helium abundance released after the M 7.3 earthquake. Errors are two sigma values. Temporal variations of helium isotopes. The location-dependent variations of 3He/*He ratios are similar before and after the 2016 Kumamoto earthquake as stated above (Fig. 2a). However, their temporal changes are different at particular sites. At Otsu hot spring, 2.2 km northwest of the Futagawa-Hinagu fault, the 3He/4He ratio decreased from 4.795 ± 0.088 Ra (all errors are 2o hereafter) in August 2010 to 4.120 ± 0.100 Ra in April 2016. Another lowering of the ratio from 1.959 ± 0.056 Ra to 1.780 ± 0.040 Ra was found at Mifune hot spring, 2.7km east of the Hinagu fault. Although Mifune is located closest to the main shock epicenter,its helium isotopes decrease is smaller than that at Otsu. This is partly due to the difference of originalHe/He ratios. If we assume that the same amount of crustal helium, 2.6 × 10-6 cm3 STP/g, was added into the Otsu and Mifune hot springs, decrease of the 3He/*He ratio becomes 0.50 Ra at Otsu, larger than 0.17 Ra at Mifune. There is no statistically valuable change of the ratio in Hirayama hot spring, 40 km away. Figure 2b shows a relationship between the distance from the fault and *He/*He changes before and soon after the earthquake. The temporal changes are relatively large in the area close to the fault and they are negligibly small at the both ends, Hirayama and Ajisai hot springs. It is, perhaps, a mirror image of the location-dependent variations (Fig. 2a), suggesting that the decrease of *He/*He ratio occurred significantly in the place where the original mantle contribution was large. Helium isotopes of rock samples. We have also measured helium abundance and isotopic ratios of six aquifer rock system of deep groundwater such as those at Mifune and Otsu (STable 2). The average He abundance of (1.9 ± 0.9) × 10-6 cm3 STP/g is nearly the same as in the Kobe aquifer rocks7 (2.1 ± 1.3) × 10-6 cm3 STP/g and is well within the variation of helium contents, (6.1 ± 4.3) × 10-6 cm3 STP/g, in crustal rocks adopted by the rock <0.01 Ra to 0.56 Ra, even though their average with a weighted error of 0.09 ± 0.02 Ra (MSWD = 0.90) is appar- ently higher than that of the Kobe samples, 0.02 ± 0.01 Ra7. Discussion s s ee sa n hypothesis that decrease of the ratio was strongly affected and perhaps invoked by the 2016 Kumamoto earth- due to diffusion, partition, and radioactive decay2. It is necesary to explain the temporal variations by physical SCIENTIFIC REPORTS|6:37939|DOl:10.1038/srep37939 3 www.nature.com/scientificreports/ means. First, we consider absolute concentrations of 4He and 2oNe in groundwater samples in the region. The 4He contents decreased at a few hot springs after the earthquake, whereas there are no meaningful changes at the other sites. There is no systematic relation with the distance and variation (SFig. la). The 2oNe contents abundances related to the M 7.3 earthquake. Second, we consider the *He/2°Ne ratios together with the 3He/*He ratios. Generally, a three-component mixing model was taken into account in subduction zones based on these ratios24. Their end members are air saturated water (ASW) with 3He/4He = 1 Ra, 4He/20Ne = 0.268; upper man- tle with 3He/*He = 8 Ra, *He/20Ne = 1000; and radiogenic with 3He/He = 0.02 Ra, *He/20Ne = 1000, respectively. Our samples together with those before the M 7.3 earthquake are located well within the mixing region of the three components in 4He/20Ne-3He/4+He diagram (SFig. 2). The 3He/4He ratios at most sites decreased after the earthquake (Fig. 2b). At the same time, the 4He/2oNe ratios also decreased. The variations cannot be attributed to simple enhancement of air contribution, because the sample with a ?He/*He ratio lower than the air value (e.g. Tamana hot spring) also shows a decrease in the 4He/2°Ne ratio. In addition, changes of a few samples (Tamana and Ueki hot springs) are not approaching into the ASW value (SFig. 2). This suggests that the temporal variation was derived from the addition of radiogenic helium together with the air component. The latter com- ponent may be attributable to the mixing of shallow and air saturated groundwater probably induced by the M 7.3 earthquake. It is possible to correct the 3He/4He ratio for diminishing air contribution based on the observed 4He/2°Ne ratio25. if the 4He/20Ne ratio is close to its air value, the correction could be significantly erroneous26. Therefore we did not take into account the Ajisai samples with low 4He/2oNe ratios of less than 0.5 in following discussion, because more than 50% of their helium is derived from air. Then we can compare the corrected 3He/*He ratios before and soon after the M 7.3 earthquake in order to estimate the increase of radiogenic helium (STable 3). Again it is possible to estimate the amount of additional crustal helium using the change ofthe corrected He/*He ratio and original abundance of helium in groundwater sample in 201018 under the assumption that 3He fluxes into hot springs are constant and helium released from aquifer rocks has a 3He/4He ratio of 0.09 Ra (STable 2). The larger addition of helium is found at Otsu and Mifune hot springs, while Tamana and Hirayama show a relatively small value. Figure 2c indicates a relationship between distance from the fault and additional helium abundance in groundwater. There is a striking trend of emanation with the distance, suggesting that the additional helium may be attributable to the M 7.3 earthquake. We should discuss the mechanism of crustal helium degassing due to the earthquake. At the time of the 1995 Kobe earthquake (M 7.2), similar degassing of helium was observed in shallow groundwater at Nishinomiya city, 30 km northeast of the epicenter'. The degassing was attributed to release of radiogenic helium accumulated in country rock as a result of micro-fracturing during the earthquake. It is possible to calculate the amount of helium released from aquifer rock into deep groundwater. Porosity of deep aquifer (Φ) was reported to be approximately 10% in the Kumamoto region27. Assuming the density of country rock (p) to be 2.8g/cm? and taking the porosity into account, additional helium in groundwater would be converted into absolute amount of helium released from the aquifer rock as follows: [He]water X Φ/{(1 - Φ)p}. This value is approximately 1.1-1.4 × 10-7 cm3 STP/g at Otsu and Mifune hot springs close to the fault, two orders of magnitude greater than that of 1.4 × 10-9 cm3STP/g during the Kobe earthquake?. The discrepancy may be due to the difference of aquifer depth (3-20 m in Kobe and 300-1000 m in Kumamoto) and distance from the earthquake fault (10 km in Kobe and 2km in Kumamoto), or volumetric strain change. On the other hand, additional helium of Hirayama hot spring at the end is consistent with zero within the experimental error range (STable 3). These geographical variations may be related to the strain change of the source zone and/or fracturing of aquifer rock affected by the M 7.3 earthquake. Coseismic volumetric strain changes at every sites are calculated using a fault model (Methods) and geodetic and B (STable 3). Strain changes are variable with time and space. For example, it is -2.04 × 10-6 for A1 fault, 1.70 × 10-6 for A2, and -1.95 × 10-6 for B in Kikuchi site, respectively. We calculate the summation of absolute values because either positive strain change (extension) or negative (compression) may contribute rock fracturing and degassing. Although Mifune is located closer to the main shock epicenter than Otsu, its total strain change is smaller than the Otsu according to the calculation results (STable 3). This may explain the smaller decrease of changes and amounts of released helium at each site (Fig. 3). This is the first data set of quantitative linkage between seismological and geochemical observations, which would be explained by a physical mechanism such as release by rock fracturing. Assuming that the amount of helium (v) released from aquifer rock is proportional to the newly exposed surface area15, we may write the relation between the volumetric stain change (△V/V) and amount of released helium (v) as: U =a × (△V/v)2/3 (1) where a is a constant and equal to k × vo k2: constant obtained by rock fracture experiment in laboratory, Uo: initial amount of helium of the aquifer rock). Least-squares fitting of data in the relationship (a dotted curve in Fig. 3) provides the value of unknown parameter as follows: a = (1.06 ± 0.13) × 10-4, where R = 0.852 and MSWD =0.84, suggesting that the ftting is statistically valid. In the fault model and/or actual feld observation during the earthquake, volumetric strain change may take a positive (extension) or negative (compression) value. On the other hand, it is difficult to conduct a rock fracturing experiment by extension mode in laboratory. Then we take into account the summation of absolute values of calculated strain changes for △V/V. Based on the esti- mated values a = 1.06 × 10-4 and k2 = 25 ± 15 from the fracturing experiment'5, the initial amount of helium of the aquifer rock is obtained to be (4.2± 2.6) × 10-6 cm3STP/g. The estimated helium abundance is gener- ally consistent with the observed value of (1.9 ± 0.9) × 10-6 cm3 STP/g in the hypothetical Kumamoto aquifer 4 www.nature.com/scientificreportsi 1.6x10 0 i P 1.2x10 MFN (cm assing 8.0x10 KKC UKI. 4.0x10 ium TMN Hel 5x10° Volumetric strain change Figure 3. Relationship between the calculated volumetric strain change and abundance of released radiogenic helium at the sampling sites. Errors are two sigma values. The dotted curve shows a least-squares fitting of equation (1). rocks within the experimental error range (STable 2) and it is comparable to the average of helium contents, (6.1 ± 4.3) x 10-6 cm3 STP/g, in crustal rocks adopted by the rock fracturing experiment15. Conclusions We found a quantitative relationship between deep groundwater helium anomaly and volumetric strain change during the 2016 Kumamoto earthquake in Southwest Japan. The correlation is plausibly explained by the gen- eration of new surface area of rock by dilatancy in the laboratory experiment, suggesting that the groundwater helium with its abundance and isotopic composition may act as effective volumetric strain gauge. Even though this phenomena should be verified in other earthquake areas, it may lead to development of a new geochemical monitoring system to detect possible strain changes prior to a hazardous earthquake in regions where adequate geophysical observation is not available. Methods The 2016 Kumamoto earthquake. The Kumamoto earthquake sequence began with a strong foreshock (M 6.5) at local time 21:26 on 14 April 2016 with a focal mechanism of strike-slip faulting at a focal depth of ~11 km. Another big foreshock (M 6.4) occurred at 00:03 on 15 April, 7km southwest of the first one (Fig. 1). Then the main shock (M 7.3) took place at 01:25 on 16 April in the vicinity of the two foreshocks. Its focal mecha- nism is also strike-slip faulting with northwest-southeast tension, and its focal depth is 10km. None of its nodal planes are consistent with the surface traces of the known Futagawa-Hinagu fault zones, which are the most fault zones and distributed at depths of 3-17 km dipping toward the northwest30. Surface fault ruptures associated with the main shock were clearly present along the Futagawa fault with a right-lateral strike-slip offset of up to 2 m31. Another offet was observed along the northernmost Hinagu fault. The main shock caused extensive dam- age to the infrastructure along the Futagawa fault with a peak ground acceleration of 1580 cm/s2, killing at least 50 inhabitants, and over 1,800 people were injured by house collapses and subsequent land slides32. About 8,300 houses were completely destroyed, some 2,600 half-damaged, and roughly 125,800 partially damaged by severe ground shaking33- Sampling and analysis. At the sampling sites, groundwater was drawn continuously from deep well by electric submersible pump for at least a couple of hours. Then a 50 cm? lead glass container with vacuum valves at both ends were connected with the faucet using a tick wall plastic tube. After approximately 10 minute of flushing of groundwater through the entire system,both valves were closed. Sample exposure to ambient air was signif- cantly minimized during the proces In the laboratory, dissolved gases were extracted by a head space method in ultra-high vacuum. A portion of gases was introduced into all metal purification and separation vacuum line, and helium and neon were refined. The *He/2°Ne ratios together with 4He and 2oNe abundances were measured by a quadrupole mass spectrometer, while the He/*He ratios were determined bya VG5400 noble gas mass spec- trometer at Atmosphere and Ocean Research Institute, the University of Tokyo34. Volumetric strain change. Coseismic volumetric strain changes at every sampling sites are calculated for a fault model35 using geodetic data reported by Geospatial Information Authority of Japan (GSI), which are estimated from ground displacements detected by interferometric synthetic aperture radar (InSAR) of satellite and global navigation satellite system (GNSs). The data were inverted for three rectangular faults with uniform 5 www.nature.com/scientificreports/ slips in an elastic half-space, A1, A2 and B, related to the 2016 Kumamoto earthquake sequence with two major foreshocks (M 6.5 and M 6.4) and the main shock (M 7.3) (SFig. 3). Then total strain changes at the sampling. sites were estimated by summation of three independent data sets (STable 3). The strongest compression of -4.01 × 10-5 was obtained at Otsu hot spring by A1, located above the fault, while a total extension of9.70 × 10-6 Was obtained at Ajisai hot spring, the southwestern end of sampling sites (Fig. 1). There is an apparent positive correlation between the estimated strain changes and amounts of released helium at each site (Fig. 3). This is the first data set of quantitative linkage between seismological and geochemical observations, which would be explained by a physical mechanism such as release by rock fracturing. There was a laboratory study of helium degassing during rock fracturing subject to uniaxial compression15. When compressional loading of a rock exceeds half of the compressive strength to destroy, the volume of the rock undergoes an inelastic increase called "dilatancy”" due to micro-cracking processes. Dilatancy may play an important role in triggering earthquakes14. When dilatancy of rock starts, new cracks are formed at grain boundaries and/or inside mineral grains. The sur- face area of the microfracture zones should increase in a rock by new cracks. Then helium in the vicinity of the newly exposed surface may be liberated from the crystalline lattice of mineral. If this is the case, there should be a positive correlation between the degassed helium and the degree of dilatancy. Based on the compression exper- iment of crustal rocks in vacuum, an equation governing empirical relationship between the residual fraction of helium in rock (R) and the dilatancy (V/V) expressed in units of volumetric strain was proposed as follows: R = 1 - k2 × (△v/v)2/3 (2) where k2 is a constant with the value of 25 ± 15 obtained by the experiment15. Following the equation, we may write the relation between the volumetric stain change and amount of helium (v) released from aquifer rock as equation (1). References 1. Holland, H. D. The Chemical Evolution of The Atmosphere and Oceans,598 pp. (Princeton Univ. Press, 1984). 2. Self S., Widdowson, M., Thordarson, T. & Jay, A. E. Volatile fluxes during flood basalt eruptions and potential effects on the global environment: A Deccan perspective. Earth Planet. Sci. Lett. 248, 518-522 (2006). 3. Black, B. A., Elkins-Tanton, L. T, Rowe, M. C. & Peate, I. U. Magnitude and consequences of volatile release from the Siberian Traps. Earth Planet. Sci. Lett.317-318, 363-373 (2012). 4. Marty, B. & Tolstikhin, I. N. CO2 fluxes from mid-ocean ridges, arcs and plumes. Chem. Geol. 145, 233-248 (1998). 5. Wallace, P Volatile i subduction zonemagmas: concentrations and fuxes basedonmelt inclusion and volcanic gas data lc. Geotherm.Res.140,217-240 (2005). 6. Irwin, W. P & Barnes, I. Tectonic Relations of Carbon Dioxide Discharges and Earthquakes. J. Geophys. Res. 85, 3115-3121 (1980). 7. Sano, Y., Takahata, N., Igarashi, G., Koizumi, N. & Sturchio, N. C. Helium degassing related to the Kobe earthquake. Chem. Geol. 150, 171-179 (1998). 8. Lee, H.et al. G. Massive and prolonged deep carbon emissions associated with continental rifting. Nature Geoscience 9, 145-149 (2016). 9. Roeloff E. A. Hydrologic Precursors to Earthquakes: A Review. PAGEOPH 126, 177-209 (1988). 10. King, C.-Y. Gas Geochemistry Applied to Earthquake Prediction An Overview. J. Geophys. Res. 91, 12269-12281 (1986). 11. Ulomoy, V. 1. & Mavashev, B. Z. Forerunners of the Tashkent earthquake, in The Tashkent Earthquake of 26 April 1966. Akad. Nauk Uzb. SSR 188-192 (1971). 12. Igarashi, G. et al. Groundwater radon anomaly before the Kobe earthquake in Japan. Science 269, 60-61 (1995). 13. Sugisaki, R. & Sugiura, T. Geochemical Indicator of Tectonic Stress Resulting in an Earthquake in Central Japan, 1984. Science 229, 1261-1262 (1985). 14. Scholz, C. H., Sykes, L. R. & Aggarwal, Y. P. Earthquake Prediction: A Physical Basis. Science 181, 803-810 (1973). 15. Honda, M., Kurita, K., Hamano, Y. & Ozima, M. Experimental studies of He and Ar degassing during rock fracturing. Earth Planet. Sci. Lett. 59, 429-436 (1982). 16. Koike, K.et al. Controls on radon emission from granite as evidenced by compression testing to failure. Geophys. J. Int. 203, 428-436 (2015). 17. Kamaya, N. et al. Overview of The 2016 Kumamoto Earthquake. JpGU Abst MIS34-P01 (2016). 18. Horiguchi, K. & Matsuda, J. Geographical distribution of He/*He ratios in north Kyushu, Japan: Geophysical implications for the 19. Sano, Y., Marty, B.& Burnard, P. Noble gases in the atmosphere. In The Noble Gases as Geochemical Tracers. Advances in Isotope Geochemistry (ed. Burnard, P) 17-31 (Springer-Verlag 2013). 21. Kennedy, B. M. et al. Mantle Fluids in the San Andreas Fault System, California. Science 278, 1278-1281 (1997). 22. Dogan, T. et al. Adjacent releases of mantle helium and soil CO from active faults: Observations from the Marmara region of the North Anatolian Fault zone, Turkey. Geochem. Geophys. Geosyst. 10, Q11009 (2009). 23. Klemperer, S. L. et al. Mantle fluids in the Karakoram fault: Helium isotope evidence. Earth Planet. Sci. Lett.366, 59-70 (2013). 24. Sano, Y. & Wakita, H. Geographical distribution of He/He ratios in Japan: Implications for arc tectonics and incipient magmatism. J. Geophys. Res. 90, 8729-8741 (1985). 25. Rison, W. & Craig, H. Helium isotopes and mantle volatiles in Loihi Seamount and Hawaiian Island basalts and xenoliths. Earth Planet. Sci. Lett.66, 407-426 (1983). 27. Shibasaki, T. Groundwater Basin and Groundwater Flow System. In Fluid Dynamics in a Deep Sedimentary Basin (Tokai Univ. Press) Pp. 109-135 (1981). 28. Yurai, H. et al. Crustal deformation of the 2016 Kumamoto Earthquake. JpGU Abst MIS34-03 (2016). 30. Shimizu, H. et al. Urgent joint seismic observation of the 2016 Kumamoto earthquake - Seismic activities and their background -. JpGU Abst MIS34-02 (2016). 31. Kumahara, Y.et al Distribution of surface rupture associated the 2016 Kumamoto earthquake and its significance. JpGU Abt MIS34-05 (2016). 32. Aoi, S. et al. Strong motion and source processes of the 2016 Kumamoto earthquake sequence. JpGU Abst MIS34-06 (2016). 34. Sano, Y. & Wakita, H. Precise Measurement of Helium Isotopes in Terrestrial Gases Bull. Chem. Soc. Japan 61, 1153-1157 (1988). 35. Okada, Y. Internal deformation due to shear and tensile faults in a half-space. Bull Seis. Soc. Am. 82, 1018-1040 (1992). SCIENTIFIC REPORTS|6:37939|DOI: 10.1038/srep37939 6 www.nature.com/scientificreports/ Acknowledgements We gratefully acknowledge K. Horiguchi and J. Matsuda for providing the information on the groundwater sampling sites. An earlier version of the paper was reviewed by K. Notsu and N. Koizumi. This study was partly supported by the Ministry of Education, Culture, Sports, Science and Technology (MEXT) of Japan, under its Earthquake and Volcano Hazards Observation and Research Program. Author Contributions Y.S. designed the study. Y.S., T.S., T.K. and T.O. collected the samples. N.T. and T.K. performed the noble gas experiments. T.S. analyzed volumetric strain change. T.O. provided geological information. Y.S., N.T., T.K. and D.Z. wrote the paper. All authors contributed to the interpretation and discussion of the data and provided comments on and input to the manuscript. Additional Information Kumamoto earthquake in Southwest Japan. Sci. Rep. 6, 37939; doi: 10.1038/srep37939 (2016). Publisher's note: Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affliations. or other third party material in this article are included in the article's Creative Commons license, unless indicated otherwise in the credit line; if the material is not included under the Creative Commons license, users will need to obtain permission from the license holder to reproduce the material. To view a copy of this license, visit http://creativecommons.org/licenses/by/4.0/ @ The Author(s) 2016 SCIENTIFIC REPORTS|6:37939|DOI:10.1038/srep37939
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Lithos 380-381(2021)105898 Contents lists available at ScienceDirect Lithos ELSEVIER journal homepage: www.elsevier.com/locate/lithos Research Article Reappraisal of the oldest high-pressure type schist in Japan: New zircon U-Pb age of the Kitomyo Schist of the Kurosegawa Belt Kazumasa Aoki d, Daniel Pastor-Galan a.b.f, Keewook Yi 8 a Graduate School of Science,Tohoku University,Sendai 980-8578,Japan CentefrthstAsiStikvrsitndi6 FacultyfScinceandhlogyeikivesitoky833a dDepasifa GradutehofItatonlReoucecincekitiversitit0 fFrontierResearch Institute for Interdisciplinary SciencesTohokuUniversity,Sendai980-0845,Japan Korea Basic Science Institute, Ochang 28119, Republic of Korea. ARTICLE INFO ＡBSＴRACT Article history: The Kitomyo Schist from Kurosegawa Belt, Shikoku, has been long considered as the oldest records of subduction Received 16 June 2020 metamorphism in Japan, based on an early 1970s K-Ar dating of white mica. The schist consists of mafic and Received in revised form 10 November 2020 pelitic layers and occurs as a tectonic block within serpentinite. Reappraisal of the schist confirmed the schist Accepted 22 November 2020 is characterized by an epidote-amphibolite peak metamorphic facies. The mafic portion is characterized by Available online 28 November 2020 zoned amphibole + epidote + chlorite + titanite ± phengite ± rutile. The presences of relict rutile surrounded by titanite and the barroisitic cores of zoned amphibole suggest a high-pressure intermediate type metamor- Keywords: Late Paleozoic phism at the metamorphic peak (P = ~0.8-1.5 GPa and T = ~500-570 °C). The presence of Mn-rich garnet Subduction metamorphism and the lack of biotite, oligoclase and paragonite also support high-pressure intermediate type metamorphism ‘Proto-Japan' that eliminate the possibility of a typical blueschist-facies metamorphism. New SHRIMP and LA-ICPMS zircon Zircon U-Pb age U-Pb geochronology on a pelitic sample show detrital grains of Mesoproterozoic and Early Paleozoic ages, sug- Kurosegawa Belt gesting a maximum deposition age for the trench-fill sediment of ~440 Ma. Also the U-Pb data confirmed ~360 Ma overgrown rims that might have formed during the subduction zone epidote-amphibolite facies meta- rather comparable to the Late Paleozoic Renge Metamorphic Rocks and their equivalents in the Kurosegawa Belt. The Devono-Carboniferous high-pressure metamorphic rocks in Japan might have been paired with their coeval batholiths along the ‘Greater South China' margin that was extensively eroded during later tectonic processes. ? 2020 Elsevier B.V. All rights reserved. 1. Introduction Wooden, & Miyamoto, 2005). The gabbroic rocks of hanging wall ophiolite, 'Oeyama Ophiolite' shows 545 ± 3 Ma zircon U-Pb age and The Japanese islands represent the longest record of active ‘Pacific- 566 ± 95 Ma Sm-Nd whole-rock isochron age (Kimura & Hayasaka, type' orogeny resulting from oceanic subduction, convergence-related 2019). This very early stage of intra-oceanic subduction zone system arc plutonism, oceanward-accretion, and landward-erosion in the has been sometimes described as 'proto-Japan' (e.g., Isozaki, 2019; world (e.g., Isozaki, 2019; Isozaki, Aoki, Nakama, & Yanai, 2010; Isozaki et al., 2010). Isozaki et al. (2010) proposed a schematic 0ceanic Maruyama, 1997). Japan likely began its Pacific-type' oceanic subduc- arc-trench cross section of 'proto-Japan' at ~520-480 Ma that delin- tion during the latest Neoproterozoic or earliest-most Paleozoic times eated ‘450 Ma (oldest) blueschist', causing misunderstanding about following the Early Paleozoic high-pressure amphibolites and schists, the age of the oldest blueschist-facies (glaucophane-schist facies) meta- jadeitites and rodingites found in the Hida-Gaien, Oeyama, Renge and morphism in Japan. Strictly speaking, Early Paleozoic blueschist-facies Kurosegawa Belts (e.g., Kunugiza et al., 2017; Maruyama & Ueda, metamorphic rocks have not yet been described in Japan (cf., 1975; Tsujimori, 2017; Tsujimori & Liou, 2004; Tsujimori, Liou, Tsujimori, 2010). The inferred two oldest localities with high-pressure (HP)-type metamorphic rocks in Japan are: (i) the Early Paleozoic kyanite- and * Corresponding author at: Tohoku University, 41 Kawauchi, Aoba-ku, Sendai, Miyagi 980-8578, Japan. paragonite-bearing epidote amphibolite in the Oeyama ultramafic E-mail address: tatsukix@tohoku.ac.jp (T. Tsujimori). body of the Oeyama Belt (~403-440 Ma Fuko-Pass Metacumulates: https://doi.org/10.1016/j.lithos.2020.105898 0024-4937/@ 2020 Elsevier B.V. Allrights reserved. S. Matsunaga, T. Tsujimori, A. Miyashita et al. Lithos380-381(2021)105898 Tsujimori, Nishina, Ishiwatari, & Itaya, 2000; Tsujimori & Liou, 2004); (a) and (i) the Early Paleozoic pelitic and mafic schists of the Kitomyo area of the Kurosegawa Belt (~402-445 Ma Kitomyo Schists: Maruyama & Ueda, 1975). Although the Fuko Pass Metacumulates bears a mineralogical indication of HP metamorphism most likely hav- V34° ing occurred in 'proto-Japan' margin, their protolith, mainly troctolitic igneous cumulates, is not typical of Pacific-type' subduction complexes. The Fuko Pass protolith geochemistry as well as the presence of spinel granulite-facies relict minerals suggest an unusually thick oceanic pla- Kitomyo teau (Tsujimori & Ishiwatari, 2002). In contrast, the Kitomyo Schist and metasedimentary rocks. However, no previous descriptions of the schist included a detailed study on index minerals (e.g., glaucophanitic amphibole and/or lawsonite) to determine its metamorphic grade. Moreover, the only geochronological data of the Kitomyo Schist is N33° white mica K-Ar dating from the early 1970s that requires reexamina- 50 km tion (e.g., Nishimura, 1998; Tsujimori & Itaya, 1999). We conducted a petrological and geochronological reappraisal of the Kitomyo Schist to Samagawa Belt Ryoke Belt understand the first generation of subduction zone metamorphic rocks within the 'proto-Japan' arc-trench system. Chichibu Belt BKurosegawa Belt PacificOc 2. Geological outline Shimanto Belt (b) ms ps Kitomyo Schist The Kurosegawa Belt is a composite geotectonic unit, a tectonic mix- ture of pre-Jurassic components of Southwest Japan (Fig. 1); various Ch spSerpentinite fragments of the pre-Jurassic geotectonic units occur as blocks or sheets Ch : Slate (Chichibu Belt) within a serpentinite-matrix melange. The Kurosegawa Belt of Shikoku KTM08 is composed of variable scale serpentinite bodies with Late Paleozoic blueschists (equivalent of the blueschists in the Hida-Gaien, and ps Renge Belts of the Hida and Chugoku Mountains), Early Paleozoic non- metamorphosed sedimentary rocks, and Early Paleozoic granitoids and rare granulite (e.g., Aoki, Isozaki, Yamamoto, Sakata, & Hirata, 450 m 500m 2015; Hada, Ishi, Landis, Aitchison, & Yoshikura, 2001; Maruyama, 1981; Maruyama, Banno, Matsuda, & Nakajima, 1984). The Kitomyo area is located at the eastern part of the Kurosegawa ms Belt of Shikoku (Fig. 1). In this area, a fault-bounded serpentinite body KTM11 (2.5 × 4.5 km) bears amphibolite (or mafic schist) and pelitic schists. The area is about 100 km east of the Ino Formation where pelitic schists sp Sp associated with glaucophane- and barroisite-bearing schists yields 50 m、 phengite K-Ar ages of 394-352 Ma (four samples) and 327-317 Ma (two samples) (Ueda, Nakajima, Matsuoka, & Maruyama, 1980). Ac- 33°50°30 cording to Maruyama and Ueda (1975), the serpentinite contains bru- Fig. 1. (a) Simplified geological map of Shikoku delineating the different belts and the cite and anthophyllite. Both mafic and pelitic schist are characterized location of Kitomyo area, where samples were collected. (b) Detailed geological map of by the occurrence of porphyroblastic albite. Maruyama and Ueda the sampling area (ms: mafic schist. ps: pelitic schist). (1975) dated white mica (K-Ar) from two psammitic schists (sample 73040305 and 71071401) in Tohoku University, obtaining 445 Ma and 402 Ma, respectively. Based on the age, they considered that the timing package uses lithium metaborate/tetraborate fusion with inductively of metamorphism was older than 445 Ma. ductively coupled plasma mass spectroscopy (FUS-ICPMS) for the 3. Methods major- and trace-element analyses, respectively. 3.1. Petrography 3.2. Geochronology We selected two samples of the Kitomyo Schist (KTM08 pelitic schist, and KTM11 mafic schist). Textures of polished petrographic We crushed the samples with a Yasui Kikai Multi Rock Pressure and thin-sections were observed using a JEOL JSM-7001F field emission- then sieved them using Nichika Nylon Mesh (#150 [~100 μm] and #100 scanning electron microscope (FE-SEM), equipped with an EDs, Oxford [~150 μm]) to obtain the proper grain-size for concentrating zircons and INCA X-act energy dispersive X-ray spectrometers at Tohoku University. phengites. Zircons were concentrated by combining conventional mag- Major-element quantitative analyses were conducted using a 15 kV ac- netic and heavy liquid methods. Hand-picked zircon grains under a bin- celeration voltage, a 1.4 nA beam current, and a 70 s integration time in the EDS system. Specifix-40) discs and polished to expose their cores. For the polishing, We also analyzed the whole-rock composition of the mafic schist (KTM11) to constrain the nature of its protolith and perform a P-T were used. pseudosection model. The analysis was carried out at Activation Labora- Cathodoluminescence (CL) images of zircon in polished mount of tories Ltd., Canada, using Code 4Litho Lithogeochemistry Package; the zircon from two sample rocks were observed using a Hitachi S-3400N S. Matsunogo, T. Tsujimori, A. Miyashita et al. Lithos 380381 (2021) 105898 SEM, equipped with a Gatan model MiniCL system in Tohoku University. show dlistinct prograde chemical zoning in spessartine decrease toward The CL observation was conducted using a 25 kV accelerating voltage the rims. Phengite is lepidoblastic (0.31 mm in size) in the matrix and and a 90 nA probe current. has occasionally intergrown with secondary chlorite; it has a composi- In-situ zircon UPb dating was carried out in the Okayama Univer- tion with 3.33.5 Si atoms per formula unit (a.p.f.u.) for O = 11. Al- sity of Science by using a Thermo Fisher Scientific iCAP-RQ single- though some phengites in quartz-rich layers are coarse-grained, collector quadrupole ICPMS coupled to a Teledyne Cetac Technologies compositional dlifference among grain size was not confirmed. Epidlote Analyte G2 ArF excimer laser ablation (LA) system equipped with a occurs as discrete subhedral grains in association with quartz. The pres- HelEx 2 volume sample chamber. The laser ablation of zircons was con- ence of gamet + albite + phengite and the lack of biotite, oligoclase and ducted at the condition of laser spot size of 25 μum with fluence of paragonite suggest that the schist underwent a HP intermediate-type 1.8 J·cm² and repetition rate of 5 Hz. Other conditions of LA-ICPMS metamorphism rather than jadeite-glaucophane type (Miyashiro, method are referred to Aoki, Aoki, Tsuchiya, and Kato (2019) and 1961). In fact, the mineral assemblage is similar to the garnet zone of Aoki, Aoki, Tsujimori, Sakata, and Tsuchiya (2020). Zircons were also the Sambagawa metamorphic belt. dated using a SHRIMP Ile/MC instrument at the Korea Basic Science In- stitute (KBSl) Ochang Center, Korea. Analytical protocols are followed Williams (1998). and reduction of the raw data was undertaken using 4.1.2. Mafic schist (KTM11) the software *SQUID' (Lucdwig, 2001). KAr age of the phengite sepa- The sample KTM11(Figs. 2c,d) is a well-deformed, amphibolitic rates was determined in the Hiruzen Institute for Geology and Chronol- schist that consists mainly of clinoamphiboles and minor amount of ep- ogy Co.Ltd:; the analytical protocol was followed by Nagao,Nishido, idote ([Fe3+/(Fe3++Al)] = 0.280.35), titanite,rutile,phengite and Itaya, and Ogata (1984) and Itaya et al. (1991). chlorite. Foliation is defined by a preferred orientation of fine-grained acicular pale-greenish actinolite and minor lepidoblastic phengite 4.Results (3.53.6 Si a.p.f.u.). Some coarse-grained blue-greenish barroisitic am- phiboles (0.51 mm in size; IBlNa [Na in the B-site] values reach up to 4.1. Mineralogical and petrological characteristics 0.67: Fig. 3) are wrapped around by layers of acicular actinolite (Figs. 2c and 3). Such textural relations indicate that relict, coarser- 4.1.1. Pelitic schist (KTM08) grained, blue-greenish barroisitic amphibole underwent grain-size re- Sample KTM08 (Figs. 2a,b) is a quartzo-feldspathic mica schist with duction by recrystallization during deformation. Some titanites reach quartz-rich layers. Porphyroblastic albite is scattered in the matrix, and up to 1 mm in size and contains abundant rutile (Fig. 2d); such textural it consists mainly of quartz, albite, phengite, secondary chlorite, and relations indicate that relict rutile was replaced by titanite during retro- minor amount of epidote ([Fe3+/ (Fe3+ + Al)] = 0.190.24), tourma- grade metamorphism. The occurrence of relict rutile together with line, ilmenite, and apatite. Oriented lepidoblastic phengite defines a barroisitic amphibole and epidote supports a HP intermediate type penetrative schistosity. Most garnets (Fig. 2b) (aalm37-sggrs2333- metamorphism. Moderate to high IBINa content of amphibole is indica- sspS13Ppyr<1) are very small (0.030.08 mm) subhedral to euhedral tive of high-pressure (e.g., Hosotani & Banno, 1986; Nakamura & Enami, grains and included within porphyroblastic albite. Euhedral grains 1994; Okamoto & Toriumi, 2005; Otsuki & Banno, 1990). (a) Oeyama/Renge &Kurosegawa Gr Ab Paleozoic and older Japan. Phe Ab South Kitakami - Kurosegawa Phe 0.2mm 50um South China craton Rt Kamuikotan/Sanbagawa/Tokoro Accretionary complexes Act Ttn Bar Act Act+Chl 0.2mm 0.2mm Ttn Fig. 2. Microphotographs showing the mineral assemblageof the collected samples. (a) Photomicrograph ofthe KMTO8 pelitic schist, showing albite,garnet and phengite. (b) FE-SEM-EDS X-ray (Mn) image of small garnets in porphyroblastic albite. (c) Photomicrograph ofthe KTM11 maficschist showing barroisite, titanite rutile and actinolite. (d) Photomicrograph showing the rutile replaced by titanite. S. Matsunogo, T. Tsujimori, A. Miyashita et al. Lithos 380381 (2021) 105898 2.0- Kitomyo, KYM 11 (a) Kitomyo mafic schist,KTM11 Sambagawa 102 Fuko Pass metacumulates [FPM] Yatsushiromafic granulites[YSGr] Tsujimori and Ishiwatari (2002) Renge Osanai et al. (2000) Type 1 Saijo gabbros[SJGb] 1.5 101 Kimura and Hayasaka (2019) 100 1.0 normali: 10-1 HOW-N 0.5 10-2, Rb C 0.0 + + (b) 0.0 0.5 1.0 1.5 2.0 [4]Al a.p.f.u.(O=23) 0.8 KTM11 0.6 Fig. 3. INa (Nainthe B-site) versus I4AI (Alin tetrahedral site) diagram for the subcalcic FPM 0.4 SGr and calcic amphiboles of Kitomyo Schist (sample KTM11). For comparisons, subcalcic and "0!1 calcic amphiboles of Sambagawa mafic schist (Okamoto & Toriumi, 2005) and sodic, subcalcic and calcic amphiboles of Renge metamorphic rocks (Tsujimori, Liou, Enst, & 0.2 SJGb SiO,wt% Itaya, 2006) are also shown. KTM11 The sample KTM11 is characterized by somewhat peculiar bulk-rock 10 FPM composition (Table 2). It is quartz normative and shows moderate SiO (51.8 wt%), low CaO (5.3 wt%) and AlO3 (9.6 wt%), with MgO (11.3 wt YSGr %),FeOT (12.2 wt%; total Fe as FeO). NazO (1.8 wt%), KO (1.09 wt%), TiO2 (0.78 wt%); the loss of ignition was 4.7 wt%. The high MgO + FeOT FeOT and Ni (450 μg·g), Cr (830 μg·g) and low CaO + AlzO and Sr SJGb SiO,wt% (31 μg·g1) (Table 2; Fig. 4) suggest that plagioclase-poor cumulate 40 45 50 55 protolith; this is also supported by a clear negative Eu anomaly in N- MORB-normalized trace-element pattern (Fig. 4a). Comparing with Fig, 4. Bulk-rock compositions of the KTM11 mafic schist sample of the Kitomyo Schist. For geochemical features with Fuko Pass metacumulates (Tsujimori & comparisons, Fuko Pass metacumulates [FPM] (Tsujimori & Ishiwatari, 2002), Yatsushiro Ishiwatari, 2002), gabbroic rocks of the Sanjo ultramafic body of the mafic granulites [YSGr] of the Kurosegawa Belt (Osanai et al., 2014) and Saijo gabbros Oeyama Ophiolite (Kimura & Hayasaka, 2019), and Yatsushiro mafic [SJGb] of the Oeyama Ophiolite (Kimura & Hayasaka, 2019) are also plotted. (a) N- MORB-normalized trace-element pattern. Normalizing values are from Sun and granulites of the Kurosegawa Belt of Kyushu (Osanai et al., 2000) McDonough (1989), except Sc, Cr and Ni from Pearce (1982). (b) SiO versus TiO and KTM11 has no similarity with those Early Paleozoic rocks, expecting a FeO° diagrams. similar FeOT and SiO contents (Fig. 4b). 4.1.3. PT condition of metamorphism T =500-570°C(Fig. 5a). Okamoto and Toriumi (2005) applied the Pelitic schist KTMo8 is characterized by a mineral assemblage of gar- Gibbs’ method for subcalcic amphiboles in mafic schists of the net + phengite + quartz + albite ± chlorite. The presence of garnet Sambagawa Belt of Shikoku and estimated its PT conditions. Using a and the lack of biotite, oligoclase and paragonite indicate that the schist new reference P-T condition for the Gibbs’ method (Uno, Iwamori, & e o es so pemp Toriumi, 2015). we calculated PT conditions of amphiboles from the mineral assemblage of the garnet zone of the Sambagawa metamorphic Kitomyo Schist and those from the Sambagawa Belt (Okamoto & belt (e.g, Aoki, Maruyama, Isozaki, Otoh, & Yanai, 2011; Itaya, Tsujimori, Toriumi, 2005) (Fig. 5a). The PT conditions based on the Gibbs' method &Liou,2011) for Kitomyo Schist overlaps with the PT estimates of barroisitic amphi- The P-T conditions for the HP intermediate type metamorphism boles of Sambagawa mafic schist (Fig. 5a). characterized by the assemblage barroisite + epidote + rutile can be The observed retrograde assemblage fits those of a typical constrained through the use of phase equilibria. Based on the bulk- greenschist facies. The absence of biotite and high-Si phengite suggests rock composition of KTM11, we modeled a PT pseudosection (equilib- P = ~0.6 GPa at T = 350 °C for the retrograde stage. The retrograde P-T rium phase diagram) using THERIAK-DOMINO software (de Capitani & path from epidote-amphibolite to greenschist facies (Fig. 5b) is rela- Petrakakis, 2010) to evaluate quantitatively the P-T stability field of tively common in HP intermediate type metamorphic belt, such as barroisitic amphibole. The pseudosection uses the thermodynamic Sambagawa Belt (Okamoto & Toriumi, 2005). Although the prograde dataset of Holland and Powell (1998); we adopted the solid solution PT path could not be constrained in the Kitomyo Schist, the retrograde models of minerals that used in Tsujimori and Ernst (2014). The calcu- evolution suggests that the Kitomyo Schist had a cooling history similar lated chemographic relations shows a PT space of coexistence of to the coherent unit of the Sambagawa Belt, before it had trapped as a barroisitic amphibole with epidote and rutile at P = ~0.81.5 GPa and tectonic block. Such retrograde path is also common in some Renge 4 S. Matsunaga, T. Tsujimori, A. Miyashita et al. Lithos 380-381(2021) 105898 (a) 1.5 + Chl, Ph ± Qz, Ab, HO 0.60 P-Testimatesbasedof Gibbs'method applied Rt using therference pointfUnoetal.2015 amphibole (Okamoto and Toriumi, 2005) Bar Ep Kitomyo (this study) BarEp BNa≥0.5●BNa<0.5 a Sambagawa (Recalculated using data Rt P from Okamoto and loriumi, 2005) G Gln Act Lws (0.40) Gln Bar Ep ssure, Ttn Ttn 1.0 0.30 GlnAct Ep S Ttn e 0.60 P 0.50 0.40 0.40- -Bt 8 1 Act Ep--- 0.30- Ilm Ttn Act Ep Bt 0.20 0.5 300 400 500 600 Temperature, °℃ 2.5 Jd +Ky (b) LwS-EC Ep-EC Pg Amp-EC 2.0 QZ a Jd+ GS P G AD e 1.5 essur Typel HGR e P EA 1.0 T02, GS AM GR 0.5 200 300 400 500 600 700 800 Temperature, °C Fig.5. (a) Equilibrium phase diagrams evaluating the stability field of KTM11 in greenschist-amphibolite facies. For comparisons, P-Testimates based on Gibb's method for barroisite and for the inferred metamorphic condition ofthe Type I and Type I Renge metamorphic rocks (Tsujimori, 2010); prograde and retrograde P-T paths (grey arrows) are after Tsujimori (2010), Tsujimori and Matsumoto (2006), and Shinji and Tsujimori (2019). metamorphic rocks (e.g., Kunugiza et al., 2017; Nakamizu, 1989). so a Tsujimori (2010) grouped such Renge Metamorphic Rocks as Type II, overgrown rims (for example, grains L1, S2 and S5 of Fig. 6). The rims which do not contain glaucophane and differ from the P-T trajectories show high CL intensity. A few grains do not show obvious internal zon- ing (grain L2) and/or exhibits distinct highly-luminescent inherited eclogite). Petrological features of the Kitomyo Schist suggest similarity core with mantled by faintly patchy dark-CL domain (grain S8). The zir- to the Type II Renge Metamorphic Rocks which do not contain con domains exhibiting oscillatory and strips of different CL intensity glaucophane (Tsujimori, 2010). suggests magmatic origin. In contrast, thin rims with bright CL are char- acteristic for hydrothermal/metamorphic overgrowths (e.g., Aoki et al., 4.2. Geochronology 2020). Twelve zircon grains were analyzed using SHRIMP and LA- ICPMS after textural observations. The 206pb/238U ages of oscillatory Zircons (~60-150 μm) from the sample KTM08 have stubby and zoned zircons show a cluster at ~440 Ma (weighted mean 443 ± 2 Ma euhedral morphology and show internal CL texture. Most grains have [MSWD = 1.92, n = 16]) and much older grains of Paleoproterozoic 5 S.Matsunaga,T.Tsujimori,A.Miyashitaetal. Lithos 380-381(2021)105898 KTM08 S1 S2 S3 S4 S5 461±15 14 L3 462±14 2 *463±15 2342 *450±1 2 *445±14 ±73 *1611±28 L11 KTM11 (*671±27 572±18 S6 S7 S8 S9 *473±15 100 μm Fig. 6. Cathodoluminescence (CL) images of thirteen zircon grains separated from the Kitomyo Schists for LA-ICPMS and SHRIMP analyses. Circles indicate the laser ablation spots for the grains L1, L2, L3 and L11 and the ion beam spot for the grains from S1 to S9. (a) 16 LA-ICPMS (b) 0900 9900 0900 9000 SHRIMP 0.14 207Pb/206Pb 0.100.12 2000 ±7Ma 1500 0.060.08 12 14 16 18 20 1000 5 10 15 238U/206Pb (b) 0.10 0.09 379 ± 21 Ma [n = 7] (207Pb/206Pb = 0.1034 ± 0.0018) 0.08 1200 206 0.07 1000 800° 0.06 600 ° 400 15 Overgro 0 10 15 238U/206Pb Fig.7. (a) Tera-Wasserburg concordia diagram for concordant data of zircons form the Kitomyo Schist (sample KTM08).(b) Enlarged plots for a ~ 440 Ma cluster and two overgrowth rims. S. Matsunaga, T. Tsujimori, A. Miyashita et al. Lithos 380-381(2021)105898 (206pb/207pb ages of 2.44 Ga, 2.34 Ga and 1.86 Ga) and Mesoproterozoic Paleozoic HP metamorphic rocks with robust evidence of 'cold' paleo- age (206pb/207Pb age of 1.53 Ga) (Table 3, Fig. 7a). Relatively wide por- geotherm. The occurrence of jadeitite associated with serpentinite de- tion of two overgrowth rims (S3 and S7 of Fig. 6) were dated using rived from the Paleozoic ophiolite and serpentinite melange of the SHRIMP. The rims yield 206pb/238U age of 362 ± 7 and 391 ± 5 Ma Hida-Gaien and the Oeyama Belt suggests an Early Paleozoic subduction n no ( initiation (Tsujimori, 2017; Tsujimori & Harlow, 2017). Early Paleozoic mantle of grain, the timing of rim overgrowth should be younger than kyanite- and paragonite-bearing metacumulates (Tsujimori et al., 391 Ma. If we consider discordant data between the two rim ages and 2000; Tsujimori & Liou, 2004) also support Early Paleozoic subduction a Mesoproterozoic age, the scattered trend defines an isochron line Zone metamorphism. On the other hand, no Early Paleozoic blueschist- with a lower intercept at 379 ± 21 Ma [n = 7] (Fig. 7b). facies mineral assemblage has been confirmed yet (Ichiyama, Koshiba, We could separate only one zircon grain from the KTM11 mafic Ito, & Tamura, 2020; Tsujimori, 2010; Tsujimori & Itaya, 1999; Tsujimori schist. The grain displays faintly planar banded zoning with a thin over- & Liou, 2004). grown rim of bright luminescence (L11 of Fig. 6). Three spot analyses on Does the Kitomyo Schist provide a clue for the first generation of the banded zoned domain did not show concordance. The apparent subduction zone along the ‘proto-Japan'? The youngest detrital zircons 206pb/238U ages, 671 ± 27 Ma, 572 ± 18 Ma, 473 ± 15 Ma (Table 3), in metasedimentary rocks of Pacific-type HP belt can constrain the max- may suggest Early Paleozoic formation of the protolith. imum depositional age of trench-fill sediments (e.g., Aoki et al., 2011). New phengite K-Ar age was shown in Table 4. It is noteworthy that Our zircon geochronology found its youngest cluster of detrital zircons the new data, 400.2 ± 7.9 Ma, for phengite separates (8.139 ± 0.16 wt at ~440 Ma. Early Paleozoic calc-alkaline granitoids with zircon U-Pb % K) overlaps the previous ages reported by Maruyama and Ueda ages of ~445-435 Ma are sporadically found as blocks in the (1975) and significantly older than the zircon rim ages. Kurosegawa Belt (e.g., Aitchison, Hada, Ireland, & Yoshikura, 1996; Aoki et al., 2015). Recently, a wide distribution of the Early Paleozoic 5. Discussion calc-alkaline granitoids has been confirmed in the Cathaysia Block of the South China Craton (e.g., Liu et al., 2014; Ou et al., 2019; Shu et al., 5.1. Significance within the ‘proto-Japan' scenario 2014; Wang et al., 2011). In the Kurosegawa Belt of Kyushu, a small ex- posure of garnet-bearing granulite and amphibolite yield zircon U-Pb The geological nature of Early Paleozoic subduction zone metamor- ages of 453-440 Ma (Osanai et al., 2014). These calc-alkaline magmatic phism in proto-Japan is poorly understood due to the paucity of Early activities and granulite-facies metamorphic rocks would have formed (a) Eroded —Detritalzircons- Cathaysia Block of SC NEJYIN Hida-Gaien* Oeyama Renge Gr Gb Jadeitite EA/AMP EC BS/EC/EA Hf model Zrn U-Pb Zrn U-Pb Hbl K-Ar Zrn U-Pb Ph K-Ar, Ar/Ar Kurosegawa* Gb Gr Kitomyo GR EA BS/EA ZmU-Pb Schist Zrn U-Pb Ph K-Ar, Ar/Ar 580 560 540 520 500 480 460 440 420 400 380 360 340 320 300 280Ma (b) ★ Devono-Carboniferous HP rocks 350Ma Panthalassa Panthalassa Kitomyo ((Renge) Laurussia Paleo-Tethys~ Sea. GSC Gondwana Fig. 8. (a)_Summary of geochronological data for metamorphic/ metasomatic and igneous rocks from the Hida-Gaien, Oeyama, Renge and Kurosegawa Belts of Japan (modified after Tsujimori & Harlow, 2012 and Tsujimori, 2017); additional data includes Ichiyama et al. (2020) and Yoshida et al. (2020). Names with *' represent composite geotectonic units. Age ranges of Late Paleozoic detrital zircons from sedimentary (and metasedimentary) rocks from the Cathaysia Block of South China Craton [SC] (Hu et al., 2012),the Yeongnam Massif [YN] (Cheong et al., 2015), Northeast Japan [NEJ] (Isozaki et al, 2014) are also shown. Abbreviations of minerals, rocks, and metamorphic facies: Zrn, zircon; PhPh, phengite; Hbl, hornblendic amphibole; Gb: gabbroic rocks; Gr, granitic rocks; BS, blueschist facies; EA, epidote-amphibolite facies; EC, eclogite facies, AM, amphibolite facies; GR, granulite facies. (b) Early Carboniferous plate reconstruction showing the location of South China (SC) and North China [NC] Cratons (modified after Young et al., 2019),'Greater South China' (Isozaki, 2019) and the Kitomyo Schist pressure metamorphic rocks (Tsujimori & Ernst, 2014) are also shown. 7 S.Matsunaga, T.Tsujimori, A.Miyashita t l. Lithos 380-381(2021) 105898 concomitantly to the first generation of ‘proto-Japan’ arc crust white micas (e.g., Itaya et al., 2011; Itaya & Tsujimori, 2015). Phengite (e.g., Isozaki, 2019). Considering regional geological context, we postu- K-Ar ages significantly older than zircon U-Pb ages have been well lated that ~445-435 Ma granites formed part of the arc-crust source for known in the eclogite-facies meta sedimentary rocks associated with the ~440 Ma detrital zircons in pelitic schist of the Kitomyo Schist. The meta-peridotite in the Sambagawa Belt (Itaya & Tsujimori, 2015). An ~1.5 Ga, ~1.8 Ga and ~ 2.4 Ga detrital grains in KTM08 might have de- older Cr-bearing phengite K-Ar age was also confirmed in metaso- rived from the cratonic blocks or inherited grains of the ~445-435 Ma matized ultramafic rocks of the Renge Belt (Tsujimori & Itaya, 1999). arc crust. Considering mantle materials have extreme 40Ar/36Ar ratio (e.g. Determining metamorphic ages using zircons in low-temperature Kaneoka & Takaoka, 1980), it is highly possible that the significantly metamorphic rocks is always challenging due to the limited zircon older phengite age is due to excess 4Ar derived from the surrounding growth under such conditions (e.g., Hay & Dempster, 2009). As our ultramafic rocks (e.g., Itaya & Tsujimori, 2015). Another possibility case has shown, metamorphic overgrowths are volumetrically too would be the detrital origin of white mica. If the closure temperature small for analyses. However succesful spot analyses revealed the is high as ~600 °C (e.g., Gozu, Yagi, Thanh, Itaya, & Compagnoni, 2016; timing of metamorphic overgrowths as young as ~360 Ma (Fig. 8). Con- n d sidering the inferred P-T trajectory of the Kitomyo Schist, metamorphic age. However, moderate to high-Si feature (3.3-3.6 a.p.f.u.) of the phengite of the Kitomyo Schist exclude the possibility. condition rather than greenschist-facies overprinting. We interpret the Recently Yang, Santosh, Maruyama, and Nakagawa (2016) and Hu rims ages represent a timing of epidote-amphibolite facies metamor- et al. (2017) conducted zircon U-Pb geochronology of blueschist, phism of the Kitomyo Schist, in which the assemblage barroisitic am- rodingites and host serpentinites in the Kurosegawa Belt near Kochi phibolite + rutile was stable. The ~360 Ma HP intermediate-type City. Yang et al. (2016) found 505 ± 3 Ma and 503 ± 3 Ma magmatic zir- metamorphism is coeval with the timing of HP metamorphism of the cons inherited in a low-grade blueschist; Osanai et al. (2014) has also Renge Metamorphic Rocks as defined by zircon U-Pb ages (Tsujimori, dated magmatic zircons inherited in a Late Paleozoic blueschist-facies 2010; Yoshida, Taguchi, Ueda, Horie, & Satish-Kumar, 2020). Then metagabbro and found 493 ± 4.9 Ma. Hu et al. (2017) found a wide why the age is significantly younger than new phengite K-Ar age? age range from 51 Ma to 1.58 Ga in rodingite samples: 197 ± 5 Ma (n = 5),262 ± 2 Ma (n = 4), 278 ± 9 Ma (n = 3), 294 ± 6 Ma (n = termine the cooling ages of HP metamorphic rocks. However,it has been 2),315 ± 11 Ma (n = 3), 811 ± 11 Ma (n = 4). They also confirmed also known chronological discrepancy due to the excess 4Ar trapped in a wide age range from 62 Ma to 2.53 Ga with clusters of 379 ± 15 Ma Table 1 Rappresentative SEM-EDS analyses for the Kitomyo Schist. Abreviation: Grt, garnet; Ph, phengite; Ep,epidote; Bar,barroisitic amphibole; Chl, chlorite. KTM08 KTM11 Grt (rim) Grt(core) Ph Ep Bar Bar Act Act Ep Chl Ph 37.57 37.61 SiO2 49.49 38.23 50.71 51.92 53.13 54.64 37.05 28.11 53.91 TiO2 0.09 0.27 0.35 0.16 0.11 0.21 0.01 0.88 0.07 Al203 20.27 20.5 26.95 25.4 6.68 6.87 2.70 1.07 22.13 17.87 21.85 Cr203 0.13 0.06 0.22 0.32 0.31 0.47 0.10 0.18 0.39 Fe203 10.97 15.80 FeOT 30.19 18.70 4.26 17.69 15.82 11.87 10.81 20.39 4.69 MnO 1.39 11.53 0.42 0.20 0.41 0.50 0.35 0.48 0.33 1.30 0.66 2.82 0.08 12.48 12.78 15.91 16.20 MgO 19.93 4.92 Cao 8.51 10.76 0.09 23.09 8.8 9.19 10.86 11.21 21.89 0.31 0'0 Na20 0.69 0.12 3.11 2.79 1.16 0.75 0.01 K20 10.25 0.06 0.18 0.12 0.16 0.16 10.64 Total 99.32 100.03 E0'S6 98.59 100.18 100.22 96.81 95.66 97.46 88.00 96.51 0= 12 12 11 12.5 23 23 23 23 12.5 28 11 Si 3.034 3.013 3.353 2.997 7.134 7.266 7.608 7.914 2.987 5.757 3.592 Ti 0.005 0.016 0.018 0.009 0.012 0.000 0.023 0.000 0.001 0.136 0.004 A1 1.929 1.936 2.152 2.347 1.108 1.133 0.456 0.183 2.103 4.313 1.716 0.000 0.000 0.007 0.004 0.024 0.035 0.035 0.054 0.006 0.029 0.021 Fe3+ 0.647 1.043 0.766 0.564 0.217 0.959 Fe2+ 2.039 1.253 0.241 1.038 1.085 0.857 1.092 3.492 0.261 Mn 0.095 0.782 0.000 0.028 0.024 0.049 0.061 0.043 0.033 0.057 0.000 Mg 0.156 0.079 0.285 0.009 2.617 2.666 3.396 3.498 0.000 6.084 0.489 Ca 0.736 0.924 0.007 1.940 1.326 1.378 1.666 1.740 1.891 0.068 0.002 PN 0.000 0.000 0.091 0.018 0.848 0.757 0.322 0.211 0.000 0.000 0.001 K 0.000 0.000 0.886 0.006 0.032 0.021 0.029 0.030 0.000 0.000 0.904 Total 7.996 8.003 7.038 8.006 15.207 15.156 15.017 14.980 7.979 19.936 6.989 [BINa 0.67 0.62 0.33 0.26 Mg# 0.07 0.06 0.54 0.72 0.71 0.80 0.76 0.64 0.65 Xalm 0.67 0.41 Xspsps 0.03 0.26 Xgrs 0.24 0.30 Xpyr 0.05 0.03 Fe2O = total Fe as Fe2O3; FeOT = total Fe as FeO Mg# = Mg/(Mg + Fe2+) atomic ratio S.Matsunaga, T.Tsujimori, A.Miyashita t l. Lithos 380-381(2021) 105898 Table 2 and Cocks and Torsvik, 2012; Isozaki, Aoki, Sakata, & Hirata, 2014; Bulk-rock major- and trace-element concentra- Isozaki et al., 2017; Isozaki, 2019) (Fig. 8b). Recent zircon geochronol- tions of the sample KTM11. ogy re-approved the classic idea of Isozaki (1997) that the 'proto- wt% Japan' was formed at an oceanic subduction zone between the paleo Pa- cific plate and South China Craton. Moreover Isozaki (2019) has pro- SiO2 51.75 TiO2 0.78 posed the ‘Greater South China'; this would consist of an Al203 9.63 amalgamated continental block that extended from the Yangtze plus FeOT 12.17 Cathaysia Blocks of the South China Craton, passing through Korean MnO 0.24 11.30 Peninsula to the Bureya, Jiamusi, and Khanka Blocks of Sikhote-Alin MgO Cao 5.27 (Primorye, Russia). However, the exposure of the Late Paleozoic HP Na20 1.81 metamorphic rocks (Renge metamorphic rocks and their equivalents) K20 1.09 with ~360-280 Ma is limited only to Japan, and their eastern or western P205 0.09 counterparts are missing in either Cathaysia or Sikhote-Alin. So far, all 101 4.73 described blueschists in Sikhote-Alin are younger (~250 Ma) and rather Total 98.86 similar to Suo Belt (Ishiwatari & Tsujimori, 2003). Similarly, Devono- Carboniferous batholiths do not crop out in the eastern margin of the Hg·g-1 'Greater South China'. However, there are abundant ~360-280 Ma detri- Rb 20 tal zircons in Permian and Jurassic sedimentary rocks in eastern part of 184 Th 0.78 the Cathaysia Block (Hu et al., 2012) and ~ 380-340 Ma detrital zircons U 0.29 in Permian sedimentary rocks in NE Japan (Isozaki et al., 2014). Devo- Nb 5.1 nian population (~370 Ma) of detrital zircons are also known in Ta 0.36 metasedimentary rocks of the Yeongnam Massif (Cheong, Kim, Kim, & La 5.7 12.5 Cho, 2015). These detrital signatures suggest a relatively extensive re- Ce Pr 1.68 gion of granitic magmatism existed along the 'Great South China' mar- Sr 31 gin and was subsequently eroded. Considering the information PN 7.55 together with regional geological context, geological evidence of a ma- Zr 49 Hf 1.3 ture arc-trench system during Late Paleozoic oceanic subduction is widely recorded as detrital zircons. This also suggest that during late Sm 2.33 Eu 0.313 Devonian-early Carboniferous the Greater South China terrane likely Gd 3.35 developed a paired belt (c.f. Sanbagawa Belt and Ryoke Bet (including Tb 0.59 granitic batholiths): Miyashiro, 1961; Brown, 2010). The surface erosion Dy 3.87 and further tectonic events probably erased the Devono-Carboniferous OH 0.86 Y 23.8 batholith belt that was paired with the Late Paleozoic HP metamorphic Er 2.59 rocks, such as the Kitomyo Schist. Tm 0.376 Yb 2.48 6. Conclusion Lu 0.403 IN 450 Sc 28 Reappraisal of the oldest high-pressure type schist in Japan confirmed Cr 830 that the Kitomyo Schist of the Kurosegawa Belt found that the schist is characterized by the HP intermediate-type, epidote-amphibolite facies metamorphism. The retrograde P-T path suggests that the Kitomyo Schist had a cooling history similar to the coherent unit of the Sambagawa Belt, (n = 10),467 ± 3 Ma (n = 5) and 488 ± 3 Ma (n = 10) in a before trapping as a tectonic block. The Kitomyo Schist contains ~440 Ma serpentinite sample. Although the presence of Mesozoic to Paleogene detrital magmatic zircon with very thin overgrown rims of ~360 Ma. zircons in the serpentinite of Paleozoic geotectonic unit suggests multi- Therefore, the schist is not the oldest HP type schist in Japan and rather ple hydrothermal zircon growth in serpentinite and/or neotectonic comparable to the Late Paleozoic Renge Metamorphic Rocks and their mingling of the serpentinite and younger strata. Probably more detailed equivalents in the Kurosegawa Belt. The both Kitomyo Schist and the zircon geochronology for the melange-matrix serpentinite is required Renge Metamorphic Rocks formed at the oceanic subduction zone along than that documented in previous studies. Nevertheless, some Late Pa- the 'Greater South China' margin. leozoic zircons might have related to the Late Paleozoic subduction zone Table 1 Representative SEM-EDS analyses of the major constituent metamorphism that is recorded in the Renge Metamorphic Rocks and minerals in the Kitomyo Schist. Abbreviation: Grt, garnet; Ph, phengite; their equivalents in the Kurosegawa Belt (Fig. 8a). Ep, epidote; Bar, barroisitic amphibole; Chl, chlorite. 5.2.Tectonic implications for East Asia Declaration of Competing Interest Where does the late Paleozoic oceanic subduction zone correlate The authors declare that they have no known competing financial with petrotectonic units in East Asia? Ernst, Tsujimori, Zhang, and Liou (2007） has considered the Permo-Triassic Tongbai-Dabie-Sulu- ence the work reported in this paper. Imjingang-Gyeonggi-Renge-Suo-Sikhote-Alin Orogenic Belt along the paleo-Pacific edge of cratonic Asia. The orogen is characterized by the Acknowledgments multiple events involving accretion of outboard oceanic arcs + microcontinental fragments against the East Asian margin at This research was supported by CNEAS and FRIS of Tohoku University ~320-210 Ma, including the deeply subducted sector like the Sulu- in part by grants from the MEXT/JSPS KAKENHI JP15H05212 and Dabie ultrahigh-pressure Belt. However, during the last decade growing JP18H01299 to TT and JP16F16329 to TT and DPG. This was also sup- evidence supports the 'proto-Japan' plate convergence at the eastern ported by MEXT Private University Research Branding Project (Okayama margin of the Cathaysia Blocks of the South China Craton (e.g., Cocks University of Science) and MEXT/JSPS KAKENHI JP19K04043 to KA. We S.Matsunaga, T.Tsujimori, A.Miyashita t l. Lithos 380-381(2021) 105898 Table 3 Zircon U-Pb isotopic data of zircons from the Kitomyo Schist (KTMO8 and KTM11). Spot IDs with‘L' represent LA-ICPMS data, resent discordant data In SHRIMP data, common Pb was corrected using measured 204pb. Spot ID 207Pb/235U20 206pb/238U 20 207pb/206pb 20 207pb/235U age 206pb/238U age 207pb/206pb age 20 20 20 U, μg/g Th, μg/g 232Th/238U LA-ICPMS L1#1* 0.606 0.0226 0.0713 0.00230 0.0616 0.00116 481.2 14 444.2 14 661.8 40 244 79.9 0.33 L1#2 0.549 0.0210 0.0719 0.00232 0.0554 0.00114 444.5 14 447.7 14 428.3 46 204 81.3 0.40 L1#3 0.575 0.0214 0.0744 0.00239 0.0561 0.00105 461.5 14 462.4 14 457.3 41 263 98.8 0.38 L1#4* 1.29 0.0465 0.0723 0.00233 0.130 0.00208 842.3 21 449.9 14 2094 28 208 81.5 0.39 0.592 0.00245 0.0580 472.0 19 460.5 L2#1 0.0291 0.0740 0.00211 528.5 80 49.0 15.1 0.31 L2#2* 1.02 0.0450 0.0744 0.00246 0.0995 0.00290 714.0 23 462.6 15 1614 54 48.1 14.5 0.30 L2#3* 0.635 0.0309 0.0714 0.00236 0.0645 0.00230 499.4 19 444.8 14 758.5 75 48.2 13.3 0.28 L3#1 3.10 0.130 0.237 0.00288 0.0949 0.00383 1432 32 1369 15 1526 76 453 127 0.28 L3#2* 3.95 0.166 0.252 0.00308 0.114 0.00459 1624 34 1448 16 1859 73 349 150 0.43 L3#3* 0.768 0.107 0.153 0.00305 0.0365 0.00503 578.7 61 915.2 17 7.23 0.932 0.13 L11#1* 1.29 0.080 0.110 0.00460 0.0851 0.00392 840.0 36 670.8 27 1318 89 14.8 10.5 0.71 L11#2* 0.643 0.0280 0.0762 0.00250 0.0613 0.00177 504.4 17 473.2 15 648.6 62 76.2 43.0 0.56 L11#3* 0.847 0.0441 0.0927 0.00310 0.0663 0.00265 623.3 24 571.5 18815.9 83 28.4 26.4 0.93 SHRIMP S1#1 8.85 0.2239 0.4288 0.01058 0.1497 0.00084 2323 12 2300 24 2342 4.8 846 51.9 0.061 S1#2 9.99 0.2759 0.4562 0.01191 0.1588 0.00143 2434 13 2423 26 2443 7.6 276 136 0.49 S2#1 0.532 0.0218 0.0716 0.00180 0.0538 0.00175 433.0 7.2 446 5.4 364.6 37 450 275 0.61 S2#2 0.556 0.0182 0.0729 0.00181 0.0553 0.00118 449.1 5.9 453.8 5.4 425.1 24 551 532 0.97 S3#1 0.415 0.0425 0.0578 0.00220 0.0521 0.00495 352.5 15 362.0 6.7 290.3 109 196 86.9 0.44 S4#1 0.559 0.0294 0.0696 0.00220 0.0583 0.00244 451.1 10 433.6 6.6 541.8 46 694 307 0.44 S5#1 0.547 0.0166 0.0699 0.00170 0.0567 0.00104 442.8 5.5 435.6 5.1 480.1 20 1051 552 0.52 S5#2 0.583 0.0200 0.0756 0.00188 0.0559 0.00132 466.4 6.4 470.0 5.6 448.5 26 581 281 0.48 S6#1 0.521 0.0320 0.0698 0.00193 0.0541 0.00297 425.7 11 435.0 5.8 375.5 62 162 76.4 0.47 S6#2* 0.496 0.0443 0.0716 0.00200 0.0502 0.00427 408.9 15 446.1 6.0 204.4 99 170 152 0.89 S7#1 0.480 0.0194 0.0625 0.00153 0.0557 0.00178 398.1 6.6 390.6 4.7 441.8 36 824 608 0.74 S8#1* 3.608 0.1351 0.264 0.00574 0.0993 0.00302 1551 15 1508 15 1611 28 132 81.5 0.62 S8#2* 2.561 0.0478 0.210 0.00336 0.0884 0.00085 1290 6.8 1230 0'6 1391 9.2 118 34.9 0.30 S9#1 0.531 0.0227 0.0697 0.00132 0.0552 0.00212 432.3 7.5 434.6 4.0 419.8 43 99.9 54.4 0.54 S9#2 0.531 0.0433 0.0727 0.00143 0.0530 0.00419 432.6 14 452.2 4.3 329.8 90 464 143 0.31 Table 4 de Capitani,C, Petrakakis,K, 2010.The computation of equilibrium assemblage diagrams with Theriak/Domino software. 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U-Th-Pb Geochronology by lon Microprobe: Reviews in Economic Ge- Osanai, Y, Hamamoto, , Kagami, H, Owada, M., Doyama,D.,Ando, T, 2000.Protolith and ology 7, 1-35. Yang,QY., Santosh,S., Maruyama, S., Nakagawa, M, 2016.Proto-Japan and tectonic - Sm-Nd geochronology of garnet-clinopyroxene granulite and garnet amphibolite from the Kurosegawa Belt in Kyushu, Southwest Japan. Memoirs ofGeological Society sion:evidence from zircon geochronology of blueschist and serpentinite.Lithosphere of Japan 56, 199-212. 8 (4), 386-395. https://doi.org/10.1130/L515.1. Osanai, Y, Yoshimoto, A., Nakano, N., Adachi, T, Kitano, I., Yonemura, K, Sakaki, J., Yoshida,T,TagchiT,a,HHrie,K,atisKumar, M, 0.arlCarnifrou Metamorphism in the Hida Gaien Belt, Japan: Implications for the Palaeozoic tectonic Tsuchiya, N., Ishizuka, H., 2014. LA-ICP-MS zircon U-Pb geochronology of Paleozoic granitic rocks and related igneous rocks from the Kurosegawa tectonic belt in Kyu- history of proto-Japan. Journal of Metamorphic Geology https://doi.org/10.1111/ jmg.12564. Young, A., Flamemt, N, Maloney, K., Williams, S., Matthews, K, Zahirovic, S., Muiller, RD. 43 (3), 71-99. https:/doi.org/10.2465/gkk.131126. 2019.Global kinematics of tectonic plates and subduction zones since the latePaleozoic Otsuki, M,Banno,S., 990.rograde and retograde metamrphism of hematitbearing basic schists in the Sanbagawa belt in Central Shikoku. J. Metam. Geol. 8 (4), Era. Geoscience Frontiers 10 (3), 989-1013. https://doi.org/10.1016/jgsf.2018.05.011. 425-439. https:/doi.org/10.1111/j.1525-1314.1990.tb00629.x. 11
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Available online at www.sciencedirect.com PHYSICS INTERIORS ELSEVIER Physics ofthe Earth and Planetary Interiors 152 (2005)14412 Seismic imaging of the entire arc of Tohoku and Hokkaido in Japan using P-wave, S-wave and sP depth-phase data Zhi Wang*, Dapeng Zhao GeodynamicsRes Received 3 March 2005; received in revised form 18 May 2005; accepted 3 June 2005 Abstract In order to beter understand seismic structure and seismotectonics of the entire arc of Tohoku and Hokkaido in Japan, we combined arrival time data from earthquakes beneath Tohoku and Hokkaido land areas, and beneath the Pacific Ocean to adopted 176,431 P-wave and 110,953 S-wave arrival times, from 5123 local earthquakes, and 2843 sP depth-phase data from 385 events that occurred beneath the Pacific Ocean. The 385 suboceanic events were accurately relocated by using P-wave, S-wave and s depth-phas arivalm datajontly hebtand reults cmd thmajratrs dneatdb previou sesndp forseismic coupling in the asperities tobe located around low-velocity areas onthe slab boundary under the suboceanic region. ? 2005 Elsevier B.V. Alrights reserved. Keywords: Seismic tomography, SP depth phase; Subduction zone; NE Japan and Kuril arcs 1. Introduction The Pacific plate subducts into the mantle with an angle of ~30° and at a rate of ~10 cm/year from the Japan Trench. The entire-arc region of Northeast (NE) * Corresponding author. Tel.: +81 89 927 8258; Japan (Fig. 1) is the site of important processes associ- fax: +81 89 927 8167. E-mail addresses: wang@sci.hime-u.acjp (Z. Wang), ated with the subduction of the Pacific plate,including zhao@sci.ehime-u.ac.jp (D. Zhao). serpentinization of the forearc mantle, backarc spread- 0031-9201/$ - see front matter @ 2005 Elsevier B.V. All rights reserved. doi:10.1016j.pi.205.06.010 Z. Wang, D. Zhao / Physics of the Earth and Planetary Interiors 152(2005）144-162 145 NE Japan. Few detailed offshore studies of the seismic velocity structure have been made for the entire-arc region of NE Japan and Hokkaido (Fig. 1) due to diff- culties in carrying out observations with ocean-bottom seismometers (OBS) in the Pacific Ocean. Umino et al. (1995) detected an sP depth-phase from seismograms of the suboceanic earthquakes recorded by land stations in NE Japan. Their results showed that the sP depth phase could be observed even from dis- tances of ~150km in this region. The sP depth phase is sensitive to the focal depth of the suboceanic earth- quakes. Hasegawa et al. (2000) showed that much more earthquakes occurred beneath the suboceanic region than beneath the NE Japan land area. Therefore, we could identify the sP depth phase in a large number of seismograms from the suboceanic events recorded by the dense array of high-sensitivity seismic stations in Japan. This enabled us to use the sP depth phase, Fig. 1. Map showing the tetonic framwork aroud the NE Japan together with P- and S-wave arrival times, to relocate andHokkaid(afrKamataandKdama 999)Cuvedines thtrenchesthatrprenthmajorlatoudarieEJapaand the hypocenters of the suboceanic events accurately. HokkaidThdistribonfactivandquatarylcan alongthEJpanrandKurilArcTholidretanldicat the present study region. 2. Previous studies ing and slab dehydration. Interplate earthquakes have Many previous studies have attempted to accurately occurred frequently along the subducting Pacific slab locate earthquakes that occurred beneath the fore- boundary, which has inficted widespread damage to arc region of the subduction zones. Nishizawa et al. the coastal areas of Japan through strong shaking and (1992) used arrival-time data recorded by both land tsunamis. The spatial distribution of the interplate and OBS stations to determine the focal depths of seismicity shows significant variations along the Japan shallow earthquakes near the Japan Trench. Frohlich Trench (Kawakatsu and Seno, 1983; Nishizawa et et al. (1982) determined the focal depth of earth- al., 1992; Hasegawa et al., 1994). The seismicity and quakes near the trench in the central Aleutians using seismic coupling or decoupling along the subducting the same method as that of Nishizawa et al. (1992). plate boundary might be correlated with structure het- Another method for locating earthquakes under the erogeneities. Therefore, accurate mapping of seismic suboceanic region is to use short-period teleseismic velocity variations throughout the entire-arc region of depth phases (Yoshi, 1979; Engdahl and Billington, the subduction zone will help to understand the inter- 1986) and/or long-period teleseismic depth phases plate seismic coupling, dehydration of the subducting (Herrmann, 1976; Forsyth, 1982). However, later tele- slab, and the mechanism oflarge interplate earthquakes - n u sesn ansn ie ssd p ouss in the region. Seismic velocity tomography has played atively large earthquakes that have enough energy to an important role in understanding the dynamics of generate teleseismic waves (Umino etal, 1995) There- subduction zones. Many researchers have used seismic fore, it is diffcult to apply these methods to the subo- tomography to investigate the NE Japan subduction ceanic region off NE Japan and Hokkaido. zone (e.g., Hasemi etal., 1984; Obara etal., 1986; Zhao Umino et al. (1995) used theoretical sP-P times et al., 1992a, 1994, 2002; Zhao and Hasegawa, 1993; calculated using a two-dimensional (2D) ray-tracing Hasegawa et al., 2000; Nakajima et al., 2001, 2002; method to determine the focal depths of many small Mishira et al., 2003). However, most of the previous interplate events. However, they did not determine the works have focused on studying the land areas under seismic velocity structure under the Pacific Ocean. 146 Z.Wang, D. Zhao /Physics of the Earth and Planetary Interiors 152 (2005) 144162 In their study, only the focal depths of suboceanic mic stations located in NE Japan and Hokkaido. Fig. 2 earthquakes were determined from the observed shows the hypocentral distribution of the 5123 earth- SP-P times, while the epicenters of the earthquakes quakes that occurred beneath the entire-arc region in were fixed by using the location procedure of the NE Japan and Hokaido. Two sets of data were used Observation Center for Prediction of Earthquakes and in this study. One set included P- and S-wave arrivals Volcanic Eruptions, Tohoku University. Mishira et al. from earthquakes that occurred under the land area (2003) used the same method as Umino et al. (1995) of NE Japan and Hokkaido during the period from to locate the focal depths of suboceanic events and July 1991 to August 2004. The other set included P- then determined the P-wave velocity structure under wave, S-wave and sP depth-phase arrivals from 385 the foreacrc region of NE Japan. However, they did suboceanic events that occurred from April to August not use P-wave, S-wave and sP depth-phase arrivals 2004. Table 1 shows the details of the two data sets. jointly to relocate the hypocenters of the suboceanic A total of2843 sP depth-phase arrivals were identifed earthquakes. and collected from 385 suboceanic earthquakes, which Compared with the previous studies, the present were selected from 8348 suboceanic earthquakes. The work has several advantages. Much more earthquakes ray paths of the 2843 sP depth phases are shown in and seismic stations are used. Hence, the seismic rays Fig. 5. It is found that the sP depth phase could be covered the study area much more densely and uni- observed from about 20% of the suboceanic earth- formly. Weused 176,431 P-and 110,953 S-wave arrival quakes recorded by the dense Hi-net seismic network, times, from 5123 local earthquakes recorded by 569 and that the epicentral distances of the sP depth phases seismic stations, in contrast to 92,000 P- and 52,000 observed ranged from 100 to 500 km in NE Japan S-waves from 1945 events in Mishira et al. (2003). In and Hokkaido. Only very clear sP depth phases were this work, 2843 sP depth-phase arrivals were identi- collected from the seismograms (Fig. 3). The picking fied accurately from 67,038 seismograms of the subo- accuracy of the sP depth phases and P-waves are esti- ceanic earthquakes. Based on the tomographic method mated tobe ~0.1-0.3 and ~0.1-0.2 s,respectively. The ofZhao et al. (1992a), we completed amethodto enable uncertainty of the S-wave arrival times was slightly the use of P-wave, S-wave and sP depth-phase arrivals larger than that of the P-wave arrival times. As a jointly to relocate the hypocenters of the suboceanic result, 5123 earthquakes were used for the inversion. earthquakes. Then we determined the 3D P- and S- The spatial distribution of the two groups of earth- wave velocity structures under the entire arc of Tohoku quakes is quite uniform in the study area (Fig. 2). The and Hokkaido. total number of P- and S-wave arrival times and sP depth phases is 176,431, 110,953 and 2843, respec- tively. 3. Data and method In this study, a total of 569 seismic stations were used which recorded the arrival time data from July 3.1. Data Sis stis ss (e ) toz is on 161 belonged to Hokkaido University (45), Tohoku Uni- A large number of P-wave, S-wave and sP versity (45), Hirosaki University (4),the University of depth-phase arrivals were used from shallow and Tokyo (16), Kyoto University (12), the Japan Meteoro- intermediate-depth earthquakes recorded by land seis- logical Agency (107) and the High-Sensitivity Seismic Table 1 Data sets fromtwo groups ofearthquakes usedinthis study Contents Number of arrivals Number of events Magnitude (MJMA) Depth (Km) Recorded stations SP Land earthquakes 109,207 4,738 飞 001> Suboceanic earthquakes 2,843 11,756 385 Total 2,843 176,431 110,953 5,123 4 Depth (km) Longitude Fig. 2. (a-c) Hypocentral distribution of 5123 land earthquakes and suboceanic earthquakes occurred beneath the entire arc of NE Japan and and black cross denote 385 suboceanic earthquakes and 4738land earthquakes,respectively (d The locations of the active volcanoes and seismic stations are indicated with triangles and squares, respectively. Network (332). The remaining eight stations were tem- continuities exist. We applied the method to 2843 porary stations. sP depth phases, together with 15,268 and 11,756 P- and S-wave arrivals, respectively, in order to relocate 3.2. Method the hypocenters of the 385 suboceanic earthquakes accurately. The first seismic-tomography method was devel- We adopted a grid spacing of 0.2° in the horizon- oped by Aki and Lee (1976). Subsequently, many tal direction and 10-20 km in the depth direction. In researchers improved this technique and applied it the tomographic inversion, three discontinuities (i.e., to various regions successfully (e.g., Hirahara, 1977; the Conrad, Moho and upper boundary of the sub- Thurber, 1983; Spakman and Nolet, 1988; Zhou and ducting Pacific slab) were taken into account in the Clayton, 1990; Zhao et al., 1992a, 1994, 1997a, velocity model. The three discontinuities are not simple 2002; Nakajima et al., 2001; Mishira et al., 2003; fat planes, but have complicated geometries (Horiuchi Hasegawa and Nakajima, 2004). In this work, we used et al., 1982a,b; Hasegawa et al., 1983; Zhao et al., the tomographic method developed by Zhao et al. 1992a,b, 1997a; Nakajima et al., 2001). So, it is nec- (1992a). This method is particularly useful for sub- essary to consider the three discontinuities in the ini- duction zone regions, in which various seismic dis- tial velocity model. The Conrad, Moho and upper boundary of the subducting Pacific slab determined by Zhao et al. (1992b, 1994) were adopted in the inversion. (2YWH) 3.3. Ray tracing for the sP depth phase The sP depth phase is a remarkable converted later phase that propagates as an S-wave up towards the 110 (sec) ocean floor. After reflection at the ocean floor, the S- wave is converted to P-wave, and finally reaches the stations located in NE Japan and Hokkaido. When cal- culating the ray paths and travel times of the P depth phase, a 3D perturbation algorithm (TDPA) was used to find the reflected point (also known as the bounce point) in the suboceanic region; the 3D ray-tracing algorithm of Zhao et al. (1992a) was then used to trace the ray segments from the hypocenter to the bounce point and to the station. Fig. 4 shows a schematic dia- gram of the TDPA. The topography of the ocean floor was considered when determin ing the travel time and NS ray path of the sP depth phase. The ocean floor topo- graphic data with a grid separation of ~0.03 were UD adopted based on the Interactive Mapping of Geosci- 110 (sec) entific Datasets (IGMT). The results showed that the 100 ocean floor topography had significant effects on deter- Fig. 3. Examples of the se ismograms showing the distinct sP depth mining the travel time and ray path of the sP depth phase and the focal depth. Fig. 6 shows the hypocen- suboceanic event. The hypocentral parameters are listed above the tral distributions of the 385 suboceanic earthquakes determined by using the first arrivals and the sP depth (alphabetical codes) are shown on the left of each seismogram. From the left to the right in each seismogram, three vertical lines represent phases jointly (crosses), and using the first arrivals only P-wave, sP depth phase and S-wave arrival time, respectively. (circles). NE Japan land Pacific Ocean Conrad VpVs Vp2Vs2 Moho P-wave Fig.4. of the sPdepth phase E, S, fER and RS denote epicenter, station, distance from the epicenter to the bounce point and distance from the converted point to the station, respectively. Fig. 6. 3Dhypocentral distributions of the 35 suboceanic events Fig. 5. Distribution of2843 sP depth-phase ray-paths from 385 subo- jointlywhlltead only the frst P- and S-wave arrivas. the suboceanic earthquakes and stations, respectively. is poor because of the lack of ray paths. More sP 4. Resolution and results depth-phase arrivals were identified in Hokkaido than that in Tohoku (Fig. 2); thus, the CRT resolution Fig. 7a and b shows the results ofVp and Vs checker- under the suboceanic region off Hokkaido was bet- board resolution tests (CRT) with a grid separation of ter than that beneath the suboceanic region off Tohoku 0.2° in the horizontal direction and 10-20 km in depth. (Figs. 7 and 8b and c). We assigned positive and negative velocity perturba- The final inverted results were obtained after three tions of ±3% to the grid nodes (Fig. 8a) and then iterations. The hit count at each grid node is greater calculated the travel times for this model. In addition, than 10 when we determined the velocity perturba- random errors (0.07 and 0.1 s for the P- and S-waves, tion. The P- and S-wave root-mean-square (RMS) respectively) were added tothe synthetic data set. Then travel time residuals calculated from the velocity model one-dimensional (1D) initial velocity model adapted and hypocenters after three iterations were 0.323 and from Zhao et al. (1992a) was used to invert the synthetic data. The CRT results for the P- and S-wave struc- of 10 was chosen after several tests. tures showed good resolution beneath the suboceanic Fig. 9a and b shows the plan views of the Vp and Vs region from depths of 25-60 km (Fig. 8b and c). The images at four different layers together with the distri- CRT resolution for a depth of 120 km is poor, although bution ofbackground seismicity during the period from the checkerboard pattern is well recovered. For most June 2002 to August 2004. Fig. 10a and b shows the ver- suboceanic earthquakes shallower than 10km, the tical cross-sections of the Vp and Vs images together sP depth phase could not be clearly identified, so with the distribution of active volcanoes and deep low- only a few shallow suboceanic earthquakes were used frequency micro-earthquakes (Okada and Hasegawa, (Fig. 2). Therefore, the CRT resolution at 10 km depth 2000) along the lines shown in the inserted map. Z.Wang,D.Zhao/PhysicsoftheEarthandPlanetaryInteriors152(2005)144-162 Fig. 11 shows the Vp and Vs tomography along the o d od unpns jo q rn (M≥ 7.5) (Nagai et al., 2001). et al., 1997a) together with the active volcanoes, low- The P- and S-wave velocity im frequency micro-earthquakes (Okada and Hasegawa, each other, although the Vs perturbations are larger 2000), larger historic interplate earthquakes (M≥ 6.0) than those of Vp. Low-velocity that occurred from 830 to 2001 (Zhao et al., 2002), and Longitude 142 42 Vp 3.0% Perturbation I resolution test for P-wave (a) and S-wave (b). Depth of each layer is exhibited in the upper left cor of each map. The Vs 3.0% -3.0% Fig. 7. (Continued). beneath the active volcanoes, and the low-velocity subducting slab (i.e., the southern one from 36.0°N to anomalies are imaged continuously along the NE Japan 37.0°N, the central one from 39.0°N to 41.0°N, and the and Kuril arcs, which are generally parallel to the northern one from 42.0°N to 43°N in the suboceanic down-dip direction of the subducting Pacific plate from region). the Moho discontinuity under the volcanic front to a The subducting Pacific slab is imaged clearly as depth of about 150 km. The other low-velocity zones a high Vp and Vs zone with a thickness of about are located beneath the suboceanic region above the 85 km (Fig. 10a and b), which is in good agreement Z.Wang, D.Zhao /Physics of the Earth and Planetary Interiors 152 (2005) 144-162 Japan Trench 5. Discussion 5.1. Hypocentral locations of the suboceanic earthquakes We used sP depth phase, P- and S-wave arrival times jointly to accurately relocate the 385 shallow subo- Deptl ceanic earthquakes off NE Japan and Hokaido. Fig. 6 Q·······. shows the hypocentral distributions of the 385 subo- ··· ceanic earthquakes that were determined by this study and by Hi-net. The results show that, in the horizon- tal direction, the shifts of most suboceanic earthquakes were towards the land and became larger toward the trench. In the depth direction, the relocatedhypocenters were generally shallower than those located byusing P- ： and S-wave arrivals. However, in the suboceanic region Depth (38.9-39.6°N and 143.6-144.8°E),the relocated focal depths were deeper than those determined by using P- 150 and S-wave arrivaltimes. We consider that these pat- Vp terns of the shifts in the horizontal and depth directions （b)20 have two causes: one is the upper oceanic crust with a low seismic velocity layer beneath the suboceanic region, the other is the ocean floor topography in the Pacific Ocean. By analyzing the data from the IGMT, we found that the thickness of the low-velocity layer (oceanic sediments) varied with location in the sub- oceanic region. The depth of the ocean foor, which ranges from tens to thousands of meters in the Pacific Ocean (topographic data are from the IGMT), has Vs p a q s (c)200 of waves from the suboceanic earthquakes. Therefore, we took into account the effects of the seafloor topog- 3.0% raphy when we determined the bounce points of the Fig.8. Vetical cro-setions of the checkerboad reolution tt sP depth phases, and travel times and ray paths of attern (a), P-wav P- and S-waves from the suboceanic earthquakes. The larger shift of hypocenters in the depth direction than sectionsatilcaadJaaeti that in the horizontal direction might refect the fact The velocityperturbation scale s shown at the botom. that the sP depth phase is more sensitive to the focal depth. Table 2 shows an example of the accuracy of the with the previous results (Hasegawa et al., 1994; Zhao travel time of the sP depth phase determined by using et al., 1992a, 1994). The deep low-frequency micro- the TDPA. The sP depth-phase arival times in Table 2 earthquakes with depths from 22 to 47 km are located were measured from the seismograms shown in Fig. 3, in or around the low-velocity zones beneath active which were selected from 528 seismic stations located volcanoes, which have anomalously low predominant on the land in NE Japan and Hokkaido. The seismo- frequencies of both P- and S-waves (1.0-5.5 and grams show that the sP depth-phase could be identifed 1.5-4.5 Hz, respectively) (Hasegawa and Yamamoto, accurately. The details of the hypocenter and stations 1994). are also listed in Table 2. The differences between the Z.Wang,D.Zhao/PhysicsoftheEarthandPlanetaryInteriors152(2005)144-162 153 observed and calculated travel times of the 2843 sP of the P- and S-wave velocity anc malies are similar depth phases were no larger than 0.8 s. to one another, although the amplitudes of the S-wave velocity anomalies are larger than those of the P-wave. 5.2. Seismic velocity anomalies in the crust and The images beneath the land of NE Japan are consis- mantle wedge tent with the seismic velocity tomography reported in previous studies (Zhao et al., 1992a; Nakajima et al., Fig. 9a and b shows the plan views of the P- and 2001; Tamura et al., 2002; Hasegawa and Nakajima, S-wave tomograpy together with the active volcanoes 2004). Vertical cross-sections of the P- and S-wave and earthquakes along each layer. The velocity pertur- bation was calculated from the average of the inverted In general, strong low-velocity anomalies appear in velocity values in each layer. The spatial distributions the cru and upper mantle wedge ber eath the active Vp -6(%) (a) Fig. 9. The plan views of P-wave (a) and S-wave (b) velocity images %) along four depths determined by the seismic velocity inverson. The depth of each layer is shown in the upper left corner of each map. Rea color denotes low velocity while blue color represents high velocity White circles indicate locations of earthquakes along each profle Red riangles denotethe active volcanoe along NE Japan and Kurilarcs The velocity perturbation scale is shown at the bottom. 10kn Vs-6(%) 6(%) Fig. 9. (Continued). ards the backarc side, and slow-velocity Herides (Zhou, 1990; Zhao et al., 1997b), Alaska (Zhao et al., 1995), Indonesia (Puspito et al., 1993) and the anomalies are imaged continuously along the volcanic front from North (44°N) to South (35N) (Fig. 9a and Mediterranean (Spakman et al., 1993). The results of b). The low-velocity anomalies are generally paral- these studies depicted low-velocity anomalies beneath lel to the down-dip direction of the subducting Pacific volcanoes, in the crust and mantle wedge, although the resolution scales are different from one another. It is 150 km (Fig. 10a and b). Similar structures beneath generally considered that the slow-velocity anomalies the crust and mantle wedge have been determined beneath the crust and upper mantle are caused by large by many researchers in other subduction zones, e.g., volumes of aqueous fluids that are released from the the Aletutians (Engdahl and Gubbins, 1987), Casca- subducting slab (Anderson et al., 1976). Beneath the dia (Michaelson and Weaver, 1986; Rasmussen and crust and the upper mantle, the subducting sediments Humphreys, 1988; Zhao et al., 2001), Taiwan (Roecker and the oceanic crust contain free water in the pore et al., 1987), New Zealand (Kuge and Satake, 1987; space and bound water in hydrous minerals, includ- Satake and Hashida, 1989), Tonga, Kermadec and ing talc, chlorite, brucite, lawsonite and amphibole. At Z.Wang,D.Zhao/PhysicsoftheEarthandPlanetaryInteriors152(2005)144-162 155 shallow depths, free could be expelled by the tle. Such flui might triggei r partial melting in compaction of subducted sediments and the collapse overlying mantle wedge and then reduce the seismic of porosity in the upper oceanic crust, which would velocity in the mantle wedge (Zhao et al., 1992a; decrease the seismic velocity ofthe crust and the upper Nakajima et al., 2001; Hasegawa and Nakajima, 2004), mantle (Fig. 10a and b).At greater depths, aqueous flu- 1 in the island arc (Gill, 1981). Most of the shallow i lakes dration reactions involving numerous hydrous minerals in NE Japan are distributed at depths that are shallower (Peacock, 1990; Schmidt and Poli, 1998; Hyndman than 15 km, although some occur at anomalously large and Peacock, 2003) that are associated with the dehy- dration process of the subducting Pacific slab, which (1-5.5 Hz) for both P- and S-waves. In contrast to might reflect the serpentinization of the forearc man- the normal shallowevents,thedeeplow-frequer 200 (a) scale is shown at thebottom. C-C' D-D 200 Depth E-E 100 velocity: s.High hea which areproba- xist in the vol- velocity areas in an ertically,trigger inNEJapa Shimazaki,1984,1987;Mats (Fig. 10a and b), strong heterogeneity is revealed at 1989; Tsumura et al., 1996), which clearly revealed the shallow depths above the subducting slab, in which the existence of the inclined low-attenuation Pacific omalies seem tobe caused by flu- slab, and high-attenuation bodies in the crust and the ids that are released upwards to the surface by porosity mantle wedge. This spatial distribution of the high- e or dehydration of the subducting sediments collapse attenuation bodies coincided with the low-Vp and low- and the oceanic crust. Vs anomalies identified beneath the crust and the man- Many previous geochemical, geological and geo- tle wedge in the present study. Okubo and Masunaga dynamical studies have provided evidence for the low- (1994) revealed that strong trench-parallel magne 57 Longitude Vp Vs Trench Trench M 6.06.5 7.07.5 8.0 Velocity -6(%) 0 +6(%) Fig. 11. Vp and Vs seismic images along the upper boundary of the subducting Pacifc slab. Blue color denotes high velocity whilered color indicateslow velocity. Opened circles repres Red trangles denote active volcanoes along NE Japan and Kurilarcs. Purple solidcircles represent epicenters of low frequency earthquakes (LFE) relocated by Okada and Hasegawa (2000). The contour lines show distribution of rupture areas of the great earthquakes (epicenters as stars) (M ≥ 7.5) estimated by Nagai et al 2001). The velocity perturbation and magnitude scales are shown at the botom. anomalies were ubiquitous just seaward of the vol- ing of under-thrust siliceous rocks with the overly- canic arc in the NE Japan subduction zone. In addition, ing mantle. The present results showing low seismic Peacock and Hyndman (199) suggested that imme- velocity above the slab interface might provide evi- diately above the slab nteface, talc-rich rocks migt dence for silica-rich rocks caused by the fluids ris- form in the mantle by the addition of silica trans- ing from the slab melting the talc-rich rocks in the ported by rising fluids and by the mechanical mix- mantle. Table 2 Locations of reflected points and travel times of sP depth phase determined by using the TDPA 2004-06-08, 05:11, 41.89N,145.24E, 29.96km, M4.0 Hypocentral parameters Station parameters 43.26833°N,143.43570°E, -0.00100 km AYWH SNNH 43.15783°N,143.89750°E,0.07400km SRMH 44.07883°N, 143.94930°E, -0.29500 km AYWH SNNH Contents SRMH Observed travel time (s) 74.6 87.0 Calculated travel time (s) -0.77 Residuals (s) 0.33 Latitude of converted point 41.943°N NI46'1 41.959N 145.1710E 145.186°E 145.199E Longitude of converted point -18 Elevation of converted point (km) 5.156 5.200 ERa (km) 7.260 8.370 212.860 177.810 265.060 RSb (km) a Distance from the epicenter to the bounce point. Z. Wang, D. Zhao / Physics of the Earth and Planetary Interiors 152 (2005) 144-162 5.3. Velocity heterogeneity in the suboceanic November 3, 1936, the Sanriku earthquake (M 7.5) on December 28, 1994, and the Tokachi-oki earthquake region (M 8.0) that occurred on September 26, 2003. For the Beneath the suboceanic region, three P- and S-wave subducting Pacific slab,the down-dip limit of the great slow-velocity anomalies at the upper slab boundary thrust earthquakes often corresponded to the intersec- are imaged clearly. The features of the slow-velocity tion of the thrust with the forearc mantle (Springer anomalies under the suboceanic region are consistent and Forster, 1998), which might be explained by with the previous studies of Ito et al. (2000), Zhao et aseismic hydrous minerals, such as serpentinite, al. (2002) and Mishira et al. (2003). This indicated which are present in the forearc mantle wedge and that the low-velocity zones above the subducting slab exhibit stable sliding (Peacock and Hyndman, 1999; are reliable features. The distinct zones (36.0-37.0°N, Oleskevich et al., 1999; Hyndman et al., 1997). Our 39.0-41.0°N and 42.0-43°N) in the Pacific Ocean with results clearly show that these interplate-thrust earth- low-velocity anomalies are imaged clearly, in which no quakes were located in areas with seismic veloci- great interplate earthquakes occurred. Because the CRT ties that are higher than those of their surrounding results for the P- and S-wave structures (Figs. 7 and 8b areas (Fig. 11), which might reflect the strong-coupled and c) show good resolution under the suboceanic asperities. In the subduction zone, the forearc man- region at depths of25-60 km, the low-velocity anoma- tle appears to be aseismic. Widespread slow-velocity lies beneath these regions are reliable features. The anomalies are revealed in the suboceanic region (with low-velocity anomalies are consistent with a proposed low-velocity perturbations of Vp and Vs of ~2-6 and seismic moment-release gap (Kawasaki et al., 2001). 4 6%, respectively) along the upper boundary of the Fluids widely exist in the crust and uppermost mantle Pacifc slab (Fig. 11). One possible explanation for the in the forearc regions of the subduction zone Tatsumi, spatial distribution of the larger interplate earthquakes 1989; Peacock, 1990; Iwamori, 1998; Zhao et al., is that the lateral heterogeneities, including strong cou- 2002). It is considered that the fluids are released pling sections or asperities and weak coupled or decou- from free pore water contained in the upper oceanic epunoq qes ddn ou suoe peoo soud pad crust and sediments, because ofthe increasing temper- might control the degree andspatial extent of the inter- ature and compressonal pressure, and/or are produced plate seismic coupling and the rupture nucleation of by the extensive dehydration of the subducting slab. the interplate earthquakes. The occurrence of slow and We suggest that the low-velocity zones correspond ultra-slow great interplate earthquakes that ruptured the to the weakly coupled or decoupled patches on the oceanic lithosphere from the seafloor to a depth of sev- slab boundary (Matsuzawa et al., 2002; Zhao et al., eral tens of kilometers (Kanamori, 1971; Kawasaki et 2002), which are causedby the presence offuid-related al., 2001) may be related to the fuid contents along the anomalies. slab upper boundary. 5.4. Interplate earthquakes on the slab boundary 5.5. Serpentinization of the forearc mantle Many large earthquakes occurred between the sub- The tomographic images (Figs. 10 and 11) show ducting plate and the overriding continental plate that considerable lateral heterogeneities exist on the along the interplate-thrust zone. Fig. 11 demonstrates upper boundary ofthe Pacific slab in the forearc region, that more than 95% of the great interplate earth- which might refect serpentinization of the forearc quakes (M≥ 7.0) occurred around the low-velocity mantle associated with dehydration of the subduct- zones, while only a few large interplate earthquakes ing slab. Dehydration of serpentinite might provide (M ~ 6.0-6.5) occurred in the slow-velocity zones, a fluid source to fux magmatism during the subduc- which might indicate weakly coupled parts of the sub- tion. A large volume of water is released from the ducting slab boundary. The frequent occurrences of the serpentinized forearc mantle by heating during sub- large interplate-thrust earthquakes caused great dam- duction or continent collision (Hyndman and Peacock, age to the coastal areas of Japan by strong shaking 2003). The forearc mantle is probably aseismic because and tsunamis; e.g., the Miyagi earthquake (M 7.5) on ofthe stable-sliding serpentinite hydrated by water that Z. Wang, D. Zhao / Physics of the Earth and Planetary Interiors 152 (2005)144-162 159 is expelled upward from the under-thrusting oceanic 6. Conclusions crust and sediments. Laboratory studies have indi- cated that serpentinite generally exhibits stable-sliding Shallow suboceanic earthquakes were relocated behavior (e.g., Reinen, 2000). erpentinite is believed accurately in the Pacific Ocean and the detailed 3D seis- to exhibit stable-sliding behavior at the plate subduct- mic velocity structures were determined throughout ing rate, thus impeding rupture into the forearc mantle the entire-arc region beneath NE Japan and Hokkaido (Bostock et al., 2002). The mechanism for generating for the first time. Our results provide insight into the stable-sliding behavior in a subducting slab could be relationship among seismic velocity, large interplate the serpentinization of the mantle wedge. Dehydra- earthquakes and coupling ordecoupling zonesthrough- tion reactions within the subducting plate might release out the entire arc of the Pacific subduction zone. water into the overlying forearc mantle wedge, result- Strong low-velocity anomalies beneath the active ing in the formation of serpentine minerals and pos- volcanoes were imaged clearly and continuously along sibly other hydrous minerals, such as talc and brucite the NE Japan arc, which are generally parallel to (Peacock and Hyndman, 1999). Because serpentiniza- the down-dip direction of the subducting Pacific slab tion will dramatically reduce the seismic velocity, the from the volcanic front to the coast of the Japan Sea. mechanical strength and density in the forearc man- The strong heterogeneous structures might have been tle while increase the Poisson's ratio, magnetization caused by aqueous fluids supplied from the under- and electrical conductivity (Hyndman and Peacock, lying slab meeting the hot upwelling flow at depths 2003), the serpentinized forearc mantle is probably u d usns sded pe wos- jo highly heterogeneous and some fluids might escape The cold subducting Pacific slab shows high-velocity to the surface. Such fuids will migrate into the over- anomalies that are ~3-6% higher than that of the nor- lain mantle and decrease the seismic velocity in the mal mantle. forearc mantle by the serpentinization proces. If the On the upper boundary of the Pacific slab under forearc mantle is serpentinized, the pattern of con- the forearc region, considerable lateral heterogeneities vection will be altered in the mantle wedge because were revealed, which showed a good correlation with of the weak rheology, low seismic velocity and low the spatial distribution of large interplate-thrust earth- density. The migration of mass in the ascending flow quakes. These results indicate that strong coupling will cause the upwelling of hot mantle materials, and sections or asperities, and weak coupled or decoupled consequently produce low seismic velocity, which is patches, exist along the upper boundary of the Pacific the manifestation of mantle diapers (Hasegawa et al., plate. The large interplate-thrust earthquakes occurred 2000). The low-velocity zones in the uppermost man- outside the low-velocity areas, which might be caused tle beneath the volcanic front (Fig. i0), which are by aqueous fuids released from the subducted oceanic inclined towards the back-arc side, might correspond crust and serpentinized oceanic mantle. to the ascending fow of subduction-induced convec- tion in the mantle wedge (Hasegawa et al., 1991; Zhao et al., 1992b, 1997b). Therefore, it is believed that Acknowledgements large volumes of aqueous fuids might be released aunns o u suna up Aq spremn We thank the Hi-net Data Center and J-Array Net- oceanic crust and sediments; the subduction of the work for providing the seismic waveform data, and the oceanic lithosphere willthen cool the overlying forearc P- and S-arrival time data available on the Internet. mantle such that low-temperature hydrous serpentine Some figures in thispaper wer producedusing Generic minerals are stable in the forearc mantle. In addi- Mapping Tools (GMT) software written by Wessel tion, the materials rising from the mantle, such as and Smith (1998). This work was partially supported magma or water produced by dehydration of the sub- by research grants (Kiban-B 11440134, and Kiban-A ducting slab, might trigger the deep low-frequency 17204037) from Japan Society for the Promotion of micro-earthquakes that occur in the lowermost crust Science to D. Zhao. We thank Prof. B. Kennett and two or the upper-most mantle (Hasegawa and Nakajima, anonymous reviewers for the thoughtful comments and 2004). suggestions. Ms. Shiping Zhong is deeply appreciated 160 Z. Wang, D. Zhao /Physics of the Earth and Planetary Interiors 152 (2005) 144-162 for critically reading the manuscript and for improving Hashida, T., Shimazaki, K., 1984. Determination of seismic atten- the English expressions. intensityatamthdandurical experitJPhyE 32,299-316. 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CHEMICAL GEOLOGY ISOTOPEGEOSCIENCE ELSEVIER Chemical Geology 201 (2003) 19-36 www.elsevier.com/locate/chemgeo Sr-Nd-Pb isotopic compositions of volcanic rocks around the Hishikari gold deposit, southwest Japan: implications for the contribution of a felsic lower crust Takahiro Hosonoa*, Takanori Nakanoa.1, Hiroyasu Murakamib.2 Institute of Geoscience, University of Tsukuba, Tsukuba, Ibaraki 305-8571, Japan bMetal Mining Agency of Japan, Tokiwa Bldg., 1-24-14, Toranomon, Minato-ku, Tokyo 105-0001, Japan Received 17 April 2002; accepted 3 June 2003 Abstract The Sr-Nd-Pb isotope compositions of Late Pliocene to Pleistocene volcanic rocks around the Hishikari mine in southern their source characteristics and evolution and the characteristics of magma related to gold mineralization. The Hishikari volcanic rocks (HVR) are classified into the Kurosonsan group (2.4-1.0 Ma) in the northern area and the Shishimano group (1.7-0.5 Ma) in the southern area. Each group is composed of three subgroups of andesite and one subgroup of rhyodacite in a late-stage. In the Kurosonsan group, the 87Sr/86Sr, 206Pb/204Pb, 207pb/204Pb and 208Pb/204Pb increase, while the 143Nd/144Nd decrease with decreasing age. The evolution pattern is opposite to this in the Shishimano group. On the Sr-Nd and Pb isotope diagrams, the HVR show a mixing array between two isotopically homogeneous components: a depleted component with a MORB-like composition, and an enriched component with radiogenic Pb and Sr and unradiogenic Nd. Regional comparison of isotope data from the HVR with those of rocks in and around Japan indicates that (1) the depleted component originated from low-K and high-alumina basaltic magma upwelling from the mantle, whereas (2) the enriched component was derived from rocks in lower crust of granodioritic composition, which we considered to have developed in the eastern margin of the Eurasian subcontinental lithosphere. According to a model calculation, the assimilation of the lower crust into the HVR increased with time from 51% to 77% in the Kurosonsan group, whereas decreased from 68% to 57% in the Shishimano group. We suggest that the Shishimano rhyodacite, which crystallized in the later stages of the Shishimano group and is considered to have been responsible for gold mineralization, was formed by intense fractional crystallization from andesite magma with a depleted Sr, Nd and Pb isotopic signature. ① 2003 Published by Elsevier B.V. Keywords: Volcanic rocks; Epithermal gold deposit; Sr-Nd-Pb isotopes; Lower crust; Assimilation; Fractional crystallization Waseda University, Ohkubo 3-4-1, Shinjuku, Tokyo 169-8555, Japan. Tel.: +81-3-5286-3318; fax: +81-3-5286-3491 E-mail addresses: hosono@aoni.waseda.jp (T. Hosono), nakanot@arsia.geo.tsukuba.ac.jp (T. Nakano), h-murakami@aist.go.jp (H. Murakami). Tel.: +81-298-53-4241; fax: +81-298-51-9764. 2 Present address: Institute for Geo-Resources and Environment, Geological Survey of Japan, National Institute of Advanced Industrial Science and Technology, AIST central 7, Higashi 1-1-1, Tsukuba, Ibaraki 305-8567, Japan. Tel.: +81-298-61-3939; fax: +81-298-61-3717. 0009-2541/S - see front matter @ 2003 Published by Elsevier B.V. doi:10.1016/S0009-2541(03)00205-5 20 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 1. Introduction 2. Geologic outline Intermediate to silicic volcanism is found com- The geology of southern Kyushu comprises main- monly on the back-arc side of convergent margins and ly the Shimanto Supergroup, a Jurassic to Paleogene is associated with hydrothermal deposits throughout accretionary prism of oceanic and trench-filled sedi- the world (e.g. Lipman et al., 1978; Hedenquist and ments dominated by sandstone and shale. It is Gulson, 1992; Macfarlane, 1999; Takagi et al., 1999). intruded by middle Miocene granitoids accompa- One must first understand the genesis of arc volcanic nying tin-tungsten mineralization (Ishihara, 1984), rocks, therefore, to study the associated ore mineral- and is overlain by late Miocene to Present volcanic ization. Late Miocene to Pleistocene volcanic rocks rocks, mainly andesites with minor basalts, dacites are distributed in southern Kyushu in southwest Japan and rhyolites (Fig. 1). The active arc volcanoes Mt. and are accompanied by many epithermal gold depos- Kirishima, Mt. Sakurajima and Mt. Kaimondake its (Fig. 1). The Hishikari mine is a world-class gold form a volcanic front along the Kagoshima Graben deposit that is unusually large and has an abnormally in the central region of southern Kyushu. Late high ore grade; its reserves, including the ore already Miocene to Pleistocene volcanic rocks accompanying mined, are estimated to be 5.2 Mt with an average the epithermal gold deposits are distributed in the grade of about 40 g/t Au (Izawa et al., 1990). Volcanic western part of this volcanic front. The volcanic area activity around the Hishikari mine occurred from 2.39 is divided into three provinces from north to south- to 0.51 Ma with gold mineralization (1.25-0.60 Ma) Hisatsu, Hokusatsu and Nansatsu (Fig. 1). Nagao et accompanying the later-stage rhyodacite (NEDO, al. (1995) reported that the Hisatsu and Hokusatsu 1991; Izawa et al., 1993; MITI, 1999). According to volcanic rocks erupted on terrestrial environments oxygen isotope studies, the early-stage ore-forming and constituted plateau lava with about 100-200 m fluid in the Hishikari gold-bearing vein contained thick, and pointed out that their eruption style was magmatic fluid (Matsuhisa and Aoki, 1994) or met- different from that formed stratovolcanos in the al-bearing deep fluid (Hayashi et al., 2001). As these Quaternary age. studies illustrate, the Hishikari area is an attractive The Hishikari mine is situated about 15 km field for investigation of the geochemical evolution of northwest of Mt. Kirishima (Fig. 1). Hishikari vol- arc magma and its relationship to epithermal gold canic rocks (termed hereafter HVR) occur in the mineralization. eastern part of the Hisatsu province and are divided Stable isotopes are useful to distinguish diverse on the basis of age, distribution and petrographic regions such as oceanic slab, slab sediment, mantle characteristics into two groups: the Kurosonsan wedge and crustal material from the lower to upper group in the northern area and the Shishimano group crust as sources of arc magma (Faure, 2001). Recent in the southern area (Fig. 1). Each group is sub- studies of volcanic rocks on the eastern margin of divided into four types. From oldest to youngest, the Eurasia (Tatsumoto and Nakamura, 1991; Chung et Kurosonsan group comprises Ichiyamakawa andesite al., 1995; Castillo, 1996) have associated subconti- (KA-I), Tataraishi andesite (KA-II), Kusumoto an- nental lithosphere with the so-called Dupal anomaly desite (KA-III) and Kurosonsan rhyodacite (KRD), (Hart, 1984) characterized by high 208pb/204pb. How- whereas the Shishimano group comprises Lower ever, no similar comprehensive study of the arc andesite (SA-I), Middle andesite (SA-II), Shishimano volcanic rocks widely distributed in southern Kyushu rhyodacite (SRD) and Upper andesite (SA-III). K- has been undertaken to date. We investigated the Sr-- Ar and fission-track age data on the HVR indicate Nd-Pb isotope compositions of the volcanic rocks in that they are of Pleistocene age, with the exception the Hishikari area for the following purposes: (1) to of part of the KA-I (NEDO, 1991; MITI, 1999). identify the source regions for igneous rocks in Although the contacts between adjacent rock units southern Kyushu; (2) to discuss the evolution of are unconformable, their ages overlap slightly. Adu- volcanic rocks in the Hishikari area; and (3) to laria in the ore vein has been dated from ca. 1.25- characterize the geochemical features of the volca- 0.66 Ma by the K-Ar method (Izawa et al., 1993). nism related to ore mineralization. The mineralization age is restricted within the for- 130°E 131°E 132°E 133°E 130°70'E Active volcano Hishikariarea 32° > Epithermalgolddeposit 07 BMS12 3 al. Airacaldera Kyushu Island Nansatsu 201(2003) KRD Shikoku KA-III Kagoshin KA-II KA-I 19-36 SA-III SA-II West Philippine IIASA-I Sea Basin Sumitomo Metal Mining) (1987), Izawa and Urashima (1989), NEDO (1991), Izawa et al. (1993) and MITI (1995, 1999). Arrows in the left-hand figure indicates the direction of recent subduction of the Philippine Sea plate (Seno, 1977). Line a in Fig. 7. Solid dotted line in the right-hand figure is the boundary between the Kurosonsan group and Shishimano group. 2 22 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 mation age of the SRD, which occurs in the vicinity Herzog HSM-F36 (Nakano and Ito, 1990) and then of the Hishikari deposit. These spatial and temporal were dissolved completely by HF (1.0 ml)-HClO4 relationships suggest that the presence of SRD is (0.2 ml)-HNOs (0.7 ml) in a closed system using a related to gold mineralization. teflon bomb of about 80 cc in volume (Nakano et al., The andesites of the HVR exhibit hyalopilitic and 1988). intersertal textures with euhedral to subhedral phe- We determined Sr, Nd and Pb isotopic ratios nocrysts of augite, hypersthene, plagioclase and using a Finnigan MAT 262RPQ multicollector mass magnetite, and contain accessory minerals including spectrometer with static multi-collection. Of HVR, ilmenite, pyrite and pyrrhotite. The matrix is com- we measured 57 samples for Sr isotope and 48 posed primarily of fine-grained plagioclase, pyrox- samples for Pb isotope. For Nd isotope, we deter- ene, magnetite, cryptocrystalline minerals and glass. mined 26 representative samples. The 87Sr/86Sr and Although the mineralogy of andesites is comparable, KA-IlI characteristically contains hornblende phe- and 86Sr of the NIST-SRM987 and 143Nd/144Nd of the nocrysts (0.5-1.5 mm) and SA-II is characterized by the presence of plagioclase phenocrysts with La Jolla-standard throughout this study were reverse zoning and frequently corroded form. The 0.710240 ±0.000016 (2o, n=27) and 0.511854 ± felsic rocks KRD and SRD have different mineral 0.000010 (2o, n=10), respectively. The Pb isotope assemblages and textures. KRD consists of euhedral ratios were normalized to NIST-SRM981 value to subhedral phenocrysts of augite, hypersthene, pue 296t'SI=9dtoz/9dzoz *907L'9s= 9dtoz/9d802) hornblende, biotite, plagioclase and quartz, and mi- qdtoz/9dgoz pue qdtoz/9dzoz qdtoz/9dsoz ueau oul nor amounts of magnetite, ilmenite and pyrite. The matrix has a spherulitic structure consisting mostly values of NIST-SRM981 throughoutthis study of gray to light brown glass. The quartz crystals (n=12) were 36.5250 ± 0.0168, 15.4369 ±0.0053 show remarkable undulatory extinction. In contrast, and 16.8961 ±0.0040 (2o), respectively. The error SRD exhibits hyalophitic to aphyric texture with on the standard reflects external reproducibility of euhedral to subhedral phenocrysts of hornblende, multiple analyses. Details of analytical techniques plagioclase, minor amounts of magnetite, ilmenite are given by Na et al. (1995) with respect to Sr and pyrite. The matrix is composed mostly of glass (61） ‘1e na ononsi Aq pue sdoost PN pue and shows a striking gray to gray-white fluidal and Jeon (1999) with respect to Pb isotope. For the structure. Unlike KRD, SRD does not contain py- determination of trace elements (Sr, Nd, Pb, Rb, roxene, biotite and quartz. Sm, U and Th), a Yokogawa PMS2000 inductively coupled plasma mass spectrometry was applied using indium as the internal standard. All facilities 3. Samples and analytical procedure used for analyses are equipped in the University of Tsukuba. We collected a total of 64 samples from the HVR for chemical analysis: 12 for KA-I, 13 for KA-II, 7 for KA-III, 8 for KRD, 8 for SA-I, 6 for SA-II, 6 for 4. Isotopic composition SRD and 4 for SA-III (Fig. 1). We also collected six samples of Shimanto sedimentary rocks (SSR) from The isotopic compositions of Sr, Nd and Pb for underground in the mine. These rock samples were HVR, other volcanic rocks in southern Kyushu, and all fresh under the microscope, and their detailed SSR are listed in Table 1. Since the HVR is young petrographycal, mineralogical and lithochemical (0.5-2.4 Ma), the deviation of the present 87sr/86Sr compositions were given in Hosono (2003). To understand the isotopic characteristics of HVR, we ligible compared to the variation within each HVR also sampled seven volcanic rocks (basalt, andesite PNt+/PNeti pue IS9s/ISL8 oUL ( -OI × S-I) dno1s and dacite) in southern Kyushu (Fig. 1). They were of the HVR are moderately variable and plot almost pulverized in a tungsten carbide vessel using a T. Hosono et al. / Chemical Geology 201 (2003) 19-36 23 Table 1 Sr, Nd and Pb isotopic compositions of HVR, volcanic rocks of southern Kyushu and SSR Sample no. PN Th 87Sr/ 143Nd/ eNd 208pb/ /207pb/ /206pb/ SiO2 MgO Sr Pb Rb Sm U 144Nd ISg (ud) (udd) (udd (udd) (udd (udd (udd) (%m) (%"m) 204pb 204pb 204pb Kurosonsan group Kurosonsan rhyodacite (KRD): 0.95-1.25 Ma 13S04 67.25 1.26 218 26.2 20.6 141 6.1 3.2 14.8 0.705422 0.512535 5-2.05 38.658 15.616 18.360 14S02 71.78 1.08 245 20.8 20.5 159 4.3 3.3 14.2 0.705076 38.609 15.609 18.345 14S03 72.06 0.95 162 21.5 17.6 151 4.8 2.3 14.1 0.705043 38.606 15.607 18.347 18S01 64.49 1.09 301 20.7 11.6 83 5.1 2.0 8.7 0.704930 0.512555 1.67 38.682 15.625 18.369 18S04 65.04 0.99 284 30.8 22.0 85 6.6 2.9 13.8 0.705259 0.512541 -1.94 38.694 15.630 18.373 24929 72.15 1.06 189 25.3 19.5 153 5.1 2.9 14.5 0.705009 一 38.621 15.611 18.349 一 24929 (leached) 0.705002 Average 68.79 1.07 233 24.2 18.6 129 5.3 2.8 13.4 0.705123 0.512543 -1.89 38.645 15.616 18.357 Kusumoto andesite (KA-II): 1.10-1.24 Ma 18S02-1 63.53 3.44 304 25.9 13.2 95 5.6 2.3 10.1 0.704822 38.610 15.609 18.347 27624 65.12 0.11 266 19.8 12.7 95 5.2 2.2 9.4 0.704862 0.512573 1.31 38.595 15.603 18.344 40016 64.69 2.16 279 20.8 20.6 90 4.9 2.5 10.0 0.704862 0.512582 1.12 ：38.710 15.633 18.380 40276 64.06 1.92 259 46.5 12.6 96 10.5 2.4 10.3 0.704919 0.512597 - 0.84 38.594 15.603 18.342 98MS64 61.07 2.32 312 21.2 10.8 74 5.3 1.8 7.3 0.704802 一 38.523 15.586 18.320 98MS142 64.26 2.38 297 25.8 12.0 86 5.7 2.0 8.9 0.704835 一 - 38.554 15.597 18.332 40474 62.32 2.77 0.704958 Average 63.58 2.16 286 26.7 13.7 89 6.2 2.2 9.3 0.704866 0.512584 -1.09 38.598 15.605 18.344 Tataraishi andesite (KA-II): 1.32-1.58 Ma 13S01 60.92 2.79 0.704529 0.512619 -0.41 38.494 15.581 18.317 13S07 62.14 3.24 390 19.5 13.0 71 4.6 2.0 7.5 0.704721 38.556 515.599 18.332 40270 62.17 2.37 411 24.5 13.9 104 5.7 2.4 10.6 0.704621 0.512613 -0.53 38.596 15.604 18.347 40281 63.63 1.98 389 29.5 13.1 101 6.6 2.7 10.1 0.704658 0.512599 -0.81 38.612 15.608 18.350 40392 63.17 2.22 351 27.7 16.9 105 6.2 2.4 10.4 0.704695 一 一 38.603 15.605 18.348 98MS42 60.13 2.65 415 20.9 12.8 86 5.0 2.3 9.1 0.704666 38.602 15.518 18.347 98MS67 61.56 2.23 398 30.3 14.0 94 7.2 2.3 9.9 0.704676 38.527 15.587 18.328 98MS179 63.06 2.35 349 27.4 14.4 106 6.3 2.6 10.9 0.704794 38.584 15.600 18.345 OT82 61.26 2.58 390 22.1 10.7 84 5.1 2.0 8.2 0.704745 38.612 ：15.607 18.352 OT82 (leached) 一 一 一 一 一 - 一 0.704750 一 一 一 40256 60.23 2.98 一 一 一 一 0.704688 一 一 一 40463 63.82 2.23 0.704701 991311 60.50 2.65 0.704743 991313 59.54 3.30 0.704750 Average 61.70 2.58 387 25.2 13.6 94 5.8 2.3 9.6 0.704691 0.512610 - 0.58 38.576 15.590 18.340 Ichiyamakawa andesite (KA-I): 1.78-2.39 Ma 15S08 59.68 3.17 0.704579 38.498 15.581 18.314 17S02 62.15 2.71 339 28.1 19.3 94 6.3 2.2 10.5 0.704477 0.512658 0.35 38.510 15.587 18.310 40439 60.61 3.08 409 22.4 12.0 75 5.3 1.6 7.3 0.704586 38.536 15.590 18.325 98MS36 58.87 2.19 399 26.2 13.9 83 6.1 2.0 8.0 0.704488 0.512668 0.55 38.476 15.577 18.302 98MS177 58.00 3.10 389 20.4 11.0 70 4.8 1.8 8.0 0.704477 0.512681 0.79 38.532 15.593 18.318 98MS193 58.66 3.47 354 21.9 14.9 67 5.2 2.5 10.1 0.704491 38.559 15.597 18.327 40408 59.89 3.74 0.704224 40474 58.99 3.16 一 0.704820 一 一 991315 63.57 2.62 0.704750 一 Average 60.05 3.03 378 23.8 14.2 78 5.5 2.0 8.8 0.704543 0.512669 0.56 38.518 15.587 18.316 (continued on next page) 24 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 Table 1 (continued) 143Nd/ 208pb/ 207pb/ 206pb/ Sample no. SiO2 MgO Sr PN Pb Rb Sm U Th 87Sr/ eNd 86Sr 144Nd 204pb 204pb 204pb (dd)(udd) (udd)(udd) (udd) (udd) (udd)(%m）(%m) Shishimano group Upper andesite (SA-I): 0.51-0.58 Ma 15S05 61.17 3.63 334 24.7 13.5 85 5.4 2.1 8.8 0.704605 0.512649 0.18 38.557 15.602 18.328 15S10 58.57 4.29 422 27.4 11.6 94 6.3 2.5 10.6 0.704607 0.512621 0.37 38.545 15.589 18.337 27680 60.73 3.00 317 20.3 13.2 78 4.9 2.0 8.5 0.704571 0.512652 0.23 38.508 15.583 18.318 27680 (leached) 一 一 0.704570 27682 61.13 3.19 388 23.9 13.1 79 5.6 2.1 8.4 0.704702 38.618 15.612 18.348 Average 60.40 3.53 365 24.1 12.9 84 5.5 2.2 9.1 0.704621 0.512641 0.01 38.557 15.597 18.332 Shishimano rhyodacite (SRD): 0.66-1.0 Ma 14S06 71.84 0.40 192 23.6 20.2 130 5.2 3.1 13.6 0.704676 0.512637 -0.05 38.447 15.564 18.300 14S07 72.21 0.35 178 27.4 23.3 133 6.1 3.5 15.0 0.704752 38.602 15.610 18.337 15S02 68.88 0.41 204 25.1 16.5 130 5.6 2.4 11.7 0.704708 0.512612 -0.54 38.593 15.608 18.336 15S03 74.17 0.36 230 21.5 21.7 130 4.7 2.9 13.4 0.704726 38.490 15.577 18.311 一 一 15S03 (leached) 0.704720 18S06 71.07 0.23 182 24.3 20.5 140 5.3 3.4 14.1 0.704711 0.512537 -2.01 38.559 15.599 18.328 HAK-8-107 70.30 0.24 0.704675 Average 71.41 0.33 197 24.4 20.5 133 5.4 3.0 13.6 0.704708 0.512596 0.87 38.538 15.591 18.322 Middle andesite (SA-II):0.78-0.79 Ma 14S09 63.06 2.99 333 22.5 14.6 92 4.9 2.1 9.5 0.704947 0.512561 -1.54 38.586 15.686 18.345 14S11 0.705353 0.512517 -2.41 38.633 15.613 18.362 14S12 64.31 2.56 304 25.1 15.1 94 5.9 1.9 8.7 0.704719 38.584 15.561 18.333 14S13 0.705324 38.625 15.610 18.357 27659 61.66 2.46 313 22.7 13.8 86 5.4 2.0 9.3 0.704882 0.512551 -1.74 38.620 15.608 18.356 Average 63.01 2.67 317 23.4 14.5 91 5.4 2.0 9.2 0.705045 0.512543 -1.90 38.609 15.616 18.351 Lower andesite (SA-I): 1.25-2.01 Ma 14S01b 62.03 3.12 319 20.1 11.4 82 4.8 1.9 8.6 0.704852 0.512567 -1.43 38.628 15.611 18.348 14S01c 62.91 3.13 301 24.0 12.0 88 5.6 2.3 10.4 0.704894 38.553 15.585 18.335 14S04 61.82 3.21 317 21.0 14.1 77 4.8 2.0 8.9 0.704877 38.573 15.597 18.337 14S05 62.81 3.23 353 24.7 12.8 77 6.0 1.9 8.0 0.704795 0.512593 -0.92 38.580 015.600 18.338 14S08 62.44 3.28 355 20.5 11.0 83 5.0 2.0 8.7 0.704873 0.512598 -0.83 38.567 715.593 18.335 27725 63.60 2.11 283 33.6 12.5 84 7.9 2.1 9.4 0.704921 38.567 15.597 18.340 27732 64.12 2.08 336 26.7 12.2 85 6.1 2.0 8.9 0.704923 38.504 15.577 18.320 Average 62.82 2.88 323 24.4 12.3 82 5.7 2.0 9.0 0.704876 0.512586 1.06 38.567 15.594 18.336 Volcanic rock of southern Kyushu R9 (basalt) 54.14 4.38 936 14.2 4.3 25 5.2 0.9 2.8 0.703568 0.512800 3.13 38.279 15.553 18.210 R17 (basalt) 52.05 3.97 613 22.9 3.5 16 6.9 0.5 2.4 0.703884 0.512796 3.04 38.317 15.549 18.224 R6 (andesite) 61.23 2.64 393 11.8 2.7 12 2.7 0.4 1.6 0.704548 0.512750 2.15 38.465 515.575 18.303 R27 (andesite) 60.59 2.42 307 13.1 7.3 36 3.8 1.0 3.7 0.704768 0.512645 0.09 38.542 15.589 18.340 R30 (andesite) 62.67 3.37 273 16.3 12.0 83 5.4 2.1 8.0 0.704847 0.512605 0.69 38.488 15.593 18.281 R22 (rhyolite) 70.62 0.46 249 39.3 13.4 99 9.1 2.4 10.8 0.704766 0.512663 0.46 38.508 15.581 18.342 R35 (rhyolite) 74.75 0.12 168 11.4 17.7 109 2.8 1.8 9.7 0.705355 0.512536 2.03 38.623 15.605 18.370 Average 62.29 2.48 420 18.4 8.7 54 5.1 1.3 5.6 0.704534 0.512685 0.88 38.460 15.578 18.296 Shimanto sedimentary rock (SSR) 1WL1 76.38 1.72 86 14.8 18.1 75 2.9 1.5 9.0 0.712801 38.816 15.639 18.495 1WL2 68.65 1.92 82 19.7 18.9 139 4.0 2.3 12.3 0.712062 38.769 15.624 18.493 一 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 25 Table 1 (continued) Sample no. 143Nd/ MgO 87Sr/ 208pb/ 207pb/ 206pb/ SiO2 PN Pb Rb Sm U Th eNd 144Nd 204pb 204pb 204pb (udd) (udd) (udd) (udd) (udd) (udd) (udd) (%m）(%m) Shimanto sedimentary rock (SSR) 1WL3 73.14 1.94 87 16.6 10.1 110 3.5 1.8 8.9 0.711155 0.512302 -6.59 38.798 15.622 18.526 1WL4 67.06 2.19 83 24.1 9.9 166 5.0 2.0 11.4 0.715784 0.512260 7.42 38.968 15.649 18.549 4W4 74.58 1.33 99 26.4 17.3 232 5.4 1.8 11.1 0.725411 0.512245 7.71 38.870 15.649 18.543 H204 70.96 1.17 121 19.2 16.5 120 4.5 2.3 10.6 0.709651 38.810 15.650 18.525 Average 71.80 1.71 93 20.1 15.1 140 4.2 1.9 10.5 0.714477 0.512269 -7.24 38.839 15.639 18.522 (Fig. 2). The Kurosonsan group tends to have higher the active volcanic area) plot in a similar field to the HVR on the Pb-Pb and Sr-Nd isotopic dia- to KA-II and KA-III, to KRD. In contrast, the Shi- grams (Figs. 4 and 5), indicating that the two shimano group exhibits an opposite trend, although components were regionally and temporally respon- the range of variation is smaller than in the Ku- rosonsan group. The Pb isotopic ratios of the HVR are also moderately variable and plot linearly on 0.51274(a) 40 207Pb/204pb and 0.51269 Jo sso ue nuay are pue kpim Aiea qdtoz/adgog the Philippine Sea basalts (Hickey-Vargas, 1991, 0.51264 0 1998) and oceanic basalts from the northern hemi- 0.51259 sphere regression line (NHRL; Hart, 1984) (Fig. 4), indicating the involvement of source materials 4 0.51254 2 Kurosonsan group change of Pb isotopes in HVR is not so distinct as 0.51249 80 × KRD -3 those of the Sr and Nd isotopes, the Kurosonsan group KA-III 0.51243 田 KA-II 06 KA-I -4 tends to have higher 208pb/204Pb with time, whereas the Shishimano group 0.51274(b) Shishimano group 2 shows the opposite trend (Fig. 3). KRD contains the 40 ●SA-III + SRD most radiogenic lead (except for one sample of KA- 0.51269 1 ●.50 II), whereas SRD has the least. endmember The most salient isotopic feature of the HVR is 0 DC-HVR 60 the linear configurations on the Sr-Nd-Pb isotope EC-HVR N 0.51259 -1 diagrams (Figs. 2 and 3). This suggests that two 144N 70 isotopically distinct components, a depleted compo- 0.51254 + -2 nent (DC-HVR) and an enriched component (EC- HVR), were responsible for the formation of the 0.51249 80 -3 HVR. From the array, it is clear that the DC-HVR M E06 871S'0<PN++1/PNt1 7t0L'0>IS98/ISL8 9Aey Pinogs 0.51243 EC 4 0.7040 0.7044 0.7048 0.7052 0.7056 87 Sr/ 86Sr have 87sr/86Sr>0.7055, 143Nd/144Nd<0.5125, 207pb/204pb>15.63 and 208pb/204pb>38.71. The volcanic rocks from other for the Shishimano group (b) of the HVR. Numbers indicate the percentage of the EC-HVR in the mixture between EC-HVR and areas in southern Kyushu (Hokusatsu, Hisatsu and DC-HVR. 26 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 EC (b) EC 38.72 latestageXX 9 38.68 90 15.62 38.64 P 204 羅 15.60 70 early stage 70 Kurosonsan group Kurosonsan group YX 60 60 × KRD × KRD KA-III 38.52 15.58 口 口 KA-II1 1 田 KA-II 田 田 KA-II X50 KA-1 ■ KA-I 38.48 endmember endmember @DC-HVR DC-HVR 15.56 EC-HVR EC-HVR 38.44 一 15.64 (c) EC (d) EC ]38.72 > X 90 38.68 90 15.62 X80 38.64 80 b 本 earlier stage 38.60 7pb/2 15.60 70 + 38.56 60 Shishimanogroup Shishimanogroup later stage, Q SA-II 1 SA-III 38.52 15.58 + SRD + SRD SA-II SA-Il 50 OSA-I 50， SA-I 38.48 endmember endmember DC-HVR F DC-HVR EC-HVR EC-HVR 38.44 15.56 18.30 18.32 18.34 18.36 18.38 18.30 18.32 18.34 18.36 18.38 206pb/204pb 206pb/204pb Fig. 3. 207Pb/204Pb vs. 206Pb/204Pb (left) and 208Pb/204pb vs. 206pb/204Pb (right) for the Kurosonsan group (a and b) and the Shishimano group (c and d), of the HVR. Numbers indicate the percentage of EC-HVR in the mixture between EC-HVR and DC-HVR. sible for the formation of all volcanic rocks in the HVR. According to Tatsumoto and Nakamura southern Kyushu. (1991), volcanic rocks in the Sea of Japan are isoto- pically of two types. The first type is alkaline rock (basalt to phonolite) from the Ulreung and Dog 5. Discussion islands, located at the eastern extension of the Korean continental plateau in the Sea of Japan (Fig. 1). This 5.1. Sr-Nd-Pb isotopic compositions of DC-HVR rock has an EM1-like Sr and Nd isotopic composition and EC-HVR (Zindler and Hart, 1986) and a Dupal isotopic anom- aly (Hart, 1984) with high 208pb. The other is tholei- Many studies have reconstructed the location of itic basalt and trachyte in the Yamato Basin. The southwestern Japan in the eastern area of the Asian plateau in the Sea of Japan has a continental structure continent before the opening of the Sea of Japan in the about 10 km thick, whereas the Yamato Basin has an middle Miocene (e.g. Otofuji and Matsuda, 1984; oceanic structure associated with a back-arc spreading Tamaki et al., 1992). Therefore, the isotopic compo- center (Ludwig et al., 1975; Kimura et al., 1987). The sitions of the igneous rocks, crust, and mantle in east Yamato Basin rocks plot tightly toward the unradio- Eurasia can provide key information on the source of genic end of the HVR array on the Sr-Nd isotope T.Hosono et al. / Chemical Geology 201(2003) 19-36 27 田田 38.8 UlreungandDogislands EC 38.4 (DC Philippineisland 38.0 endmember 37.6 sS Average of subducted sediments SSRAverage of Shimanto sedimentary rocks DC-HVR ECEC-HVR 15.65 EC S邮 15.60 northernTaiwan Philippine Sea 15.55 C volcanic rocks Ulreun andDogislands NHRL 15.50 15.45 15.40 × High-Mg andesite Volcanic rock of southern Kyushu ●Volcanic rock of Yamato Basin 15.35 Range of the Terrigenous sediment Yamato Basin rocks 田 Pelagic sediment ■Shimanto sedimentary rock 17.8 18.0 18.2 18.4 18.6 206pb/204pb Fig. 4. 206Pb/204Pb vs. 208Pb/204Pb and 207Pb/204Pb for HVR and their three end-components (DC-HVR, EC-HVR and SSR). Data for high-Mg andesite (Ishikawa and Nakamura, 1994; Shimoda et al., 1998), volcanic rocks of the late Miocene to the Present age in southern Kyusyu, terrigenous and pelagic sediments (Shimoda et al., 1998), and volcanic rocks from Philippine sea (Hickey-Vargas, 1991, 1998), Philippine Islands (Castillo, 1996), northern Taiwan (Chung et al., 1995) and Sea of Japan (Tatsumoto and Nakamura, 1991) are also plotted for comparison. SS (subducted sediment) and SSR represent the average isotopic ratios from the data sources (Table 2). Small squares on the mixing line connecting between least-radiogenic basalt of the Yamato Basin (Table 2) and SS indicate the mixing ratio with 10% interval. 28 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 0.513050 Philippineislands ▲ Aira silicic pumice (DC 0.512948 × High-Mg andesite 6 +Middle Miocene granitoid O Volcanic rock of southern Kyushu ●Volcanicrock of Yamatobasin 0.512845 O Terrigenous sediment northern 田 Pelagic sediment Taiwan O ■Shimanto sedimentary rock 0.512743 2 southern endmember O 灸 Ryukyu arc [ss Average of subducted sediments 0.512640 [SsR Average of Shimanto sedimentary- HVR 0 DC-HVR rocks PN3 EC-HVR 0.512537 Ulreung and -2 Dogislands /PN 0.512435 4 SS 0.512332 EM1 6 SSR 0.512230 ★andesite from north part of southwest Japan 0.512127 10 0.704 0.706 0.708 0.710 0.712 0.714 0.716 87Sr/86Sr Kakubuchi et al., 1994; Kagami et al., 1995; Shimoda et al., 1998), middle Miocene granitoids of Outer zone (Terakado et al., 1988; Shinjoe, 1997), terrigenous and pelagic sediments (Shimoda et al., 1998), five samples of Shimanto sedimentary rocks (Terakado et al., 1988), Neogene andesite with the lowest 143Nd/144Nd from the north part of southwest Japan (Terakado et al., 1997) and volcanic rocks from Philippine Islands (Castillo, 1996), northern Taiwan (Chung et al., 1995), southern Ryukyu arc (Shinjo, 1998) and Sea of Japan (Tatsumoto and Nakamura, 1991) are also plotted for comparison. diagram (Fig. 5), whereas its Pb isotope ratios vary arc regions can be formed by assimilation of base- from unradiogenic to radiogenic compositions (Fig. ment rocks and/or fractional crystallization of basic to 4). We can assume from these results that the Yamato intermediate magma (e.g. Johnson and Fridrich, Basin rocks are a suitable source for DC-HVR but 1990; Christensen and DePaolo, 1993; Graham et they are sensitive with respect to the Pb isotopic al., 1995). Therefore, it could be the case that the compositions because mantle where the basalt origi- nated is deficient in Pb, and is easily altered Pb- on the Philippine Sea plate are equivalent to the EC- isotopically by the incorporation of other Pb-rich HVR source. However, this possibility seems negli- sources such as subducted sediment and continental gible or small, because these sedimentary components materials (Miller et al., 1994; Brenan et al., 1995). deviate significantly from the HVR array on the Sr- Accordingly, we can anticipate the Pb isotopic com- Nd-Pb isotopic diagrams (Figs. 4 and 5). Instead, the position of DC-HVR to be about where the HVR and Yamato Basin rock arrays intersect in Fig. 4. ces from the Aira caldera in central southern Kyushu On the other hand, we can conclude from its plot at the enriched extension of the HVR array (Fig. radiogenic isotopic signature that EC-HVR is felsic. 5). Arakawa et al. (1998) presumed that the magma Many studies have confirmed that silicic magma in produced Aira silicic pumices formed with contami- T. Hosono et al. / Chemical Geology 201 (2003) 19-36 29 nation in the lower part of crust. Further, Terakado et Nd isotopic ratios of middle Miocene granitoids and al. (1997) reported that some Neogene mafic to high-Mg andesite in the Southwestern Japan (Fig. 5), intermediate volcanic rocks from southwest Japan UA18 e 101 (61ZIS'0 913Ae) PN+1/PNet1 M01 peq shown by the line connecting DC-HVR and HVR. 87sr/86Sr (average 0.7089) (Fig. 5), and suggested a lower crustal provenance. We also plotted the Sr and incorporation of SSR (Shinjoe, 1997; Ishihara and [ (a) 80 (b) EC 0.7060 0.7050 70 EC 0.7055 0.7045 50/ X % 40 X 80 30 0.7050 70 0.7040 0.2 0.4 0.6 0.8 1.0 60 0.7045 40 Kurosonsan group 30 endmember × KRD 20 0.7040 KA-III @ DC-HVR 10 田KA-II DC EC-HVR KA-I (c) 0.6 0.8 1.0 (d) 3 EC 0.7060 0.7050 70 EC 3 line 60 o 90 50 0.7055 0.7045 40- 80 simple1 30— 70 0.7050 60 交 50 fractional crystallization 0.7045 40 Shishimano group 30 endmember ●SA-III 20 0.7040 + SRD @ DC-HVR 10 DC SA-II (DC) EC-HVR O SA-I 0 2 3 4 0.01 0.12 0.14 0.16 0.18 0.20 1/MgO 1/SiO2 Fig. 6. 87Sr/86Sr vs. 1/MgO (let) and 1/SiO2 (right) for the Kurosonsan group (a and b) and the Shishimano group (c and d) of the HVR. (1991) and Hosono (2003). 30 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 Matsuhisa, 1999). We assume from these that EC- (Hosono, 2003). Based on the estimated 87sr/86Sr HVR corresponds to the lower crust whose origin is value for DC-HVR and EC-HVR (Table 2), we can different from the overlying Shimanto upper-middle estimate the respective concentrations of MgO (Fig. crust. 6a,c and Table 2) as 5.0 wt.% for DC-HVR and 2.0 wt.% for EC-HVR. However, the linearity becomes 5.2. Elemental compositions of DC-HVR and EC-HVR weak for SiO2 (Fig. 6b and d), Na2O and K2O (Hosono, 2003), which are mainly substituted in The reciprocal concentrations of major elements of plagioclase, quartz and residual melt. Most of the the HVR andesites (Hosono,2003) show a good phenocryst plagioclase exhibits normal zoning, indi- correlation with the Sr, Nd and Pb isotopic composi- cating that the crystallization continued until a later tions. Fig. 6 shows a linear relationship between stage. We consider that the slight deviation of SiO2, 87Sr/86Sr and 1/MgO for HVR, which indicates that Na2O and K2O from the simple binary mixing line is Mg in HVR was derived from a mixture of DC-HVR caused by fractional crystallization of magma, which and EC-HVR (Faure, 1986). A similar relationship resulted in their concentration in the residual melt. On exists among another isotopes (Nd and Pb) and the other hand, FeO, MnO and MgO would not elements (TiO2, Al2O3, FeO, MnO, CaO and P2Os) concentrate in residual melt, since they are essentially Table 2 Isotopic and elemental compositions of endcomponents and other rocks for comparison DC-HVR EC-HVR Basalta SSRb A B SiO2 (wt.%) 53.00 66.50 52.03 67.64 51.94 68.88 TiO2 1.25 0.56 1.09 0.64 0.85 0.76 Al2O3 17.20 15.20 17.55 15.85 18.77 14.62 FeO* 10.00 5.00 9.89 4.75 9.97 4.80 MnO 0.17 0.10 0.17 0.01 0.19 0.14 MgO 5.00 2.00 5.74 1.77 5.35 0.47 CaO 8.30 4.50 9.48 1.72 10.12 3.04 Na2O 2.50 3.00 3.06 3.13 2.36 4.21 K2O 1.67 3.00 0.68 3.33 0.71 2.61 P2O5 0.20 0.10 0.13 0.12 0.12 0.22 Total 99.29 99.96 99.82 99.02 100.38 99.75 DC-HVR EC-HVR Basaltc SSRd VYBe Ssf Sr (ppm) 500 250 936 93 632 175 PN 24.0 24.0 14.2 20.1 45 22 Pb 11.1 14.3 4.3 15.1 2 29 87Sr/86Sr 0.70380 0.70600 0.70357 0.71448 0.70351 0.70989 PNt+/PNgt1 0.512958 0.512380 0.512800 0.512269 0.512998 0.512367 PN3 6.20 5.50 3.13 7.24 6.98 5.33 208Pb/204pb 38.150 38.755 38.879 38.839 37.493 38.780 207pb/204pb 15.477 15.645 15.553 15.639 15.337 15.620 206pb/204pb 18.200 18.395 18.210 18.522 17.778 18.545 A= Composition of residual basaltic magma calculated by Yanagi et al. (1991). B = Composition of primary dacitic magma calculated by Yanagi et al. (1991). a Average compositions of basalt (n =4) from Hisatsu volcanic province (Nagao et al., 1999). b Average compositions of shale and sandstone (n=38) of Shimanto sedimentary rocks (Shinjoe, 1997). ° Basalt sample (sample no. R9) from Hisatsu volcanic province. d Average composition of six Shimanto sedimentary rocks. e The most primitive basalt of Yamato basin after Tatsumoto and Nakamura (1991). f Average compositions of terrigenous sediments from Nankai Trough (n= 7) and pelagic clay samples from the Philippine Sea (n= 3). Data source are Shimoda et al. (1998). * FeO represents the total iron content. T. Hosono et al. / Chemical Geology 201 (2003) 19-36 31 substituted in mafic minerals, and hence crystallized increased systematically with time, from an average of mostly in earlier stages. As a result, these elements do 50% for KA-I to 58% for KA-II and 64% for KA-III, not deviate significantly from the binary mixing line. to 75% for KRD. In the Shishimano group, the The binary mixing origin of the HVR andesites average contribution of EC-HVR, although it in- allows us to calculate the major elemental composi- creased slightly from 65% for SA-I to 67% for SA- tions of DC-HVR and EC-HVR from their Sr-Nd- II, decreased in the later stages to 59% for SRD and Pb isotopic ratios (Table 2). For the elements (SiO2, 55% for SA-III. The mixing values are slightly Na2O and K2O) that are susceptible to fractional scattered on the Pb isotopic diagram, but the average crystallization, we drew a mixing line passing along mixing value agrees well with that estimated by the their lowest concentrations (Fig. 6). Our calculation Sr-Nd isotope ratios (Fig. 3). Based on Pb isotopes, results show that the total concentration of major the contribution of EC-HVR increased with time for elements for both DC-HVR and EC-HVR are nearly the Kurosonsan group (53% to 66% to 69% to 75%), 100%, indicating that the binary mixing model is whereas it tended to decrease in the later stages for the satisfactory. Likewise, we calculated the Sr, Nd and Shishimano group (65% to 71% to 56% to 61%). The Pb concentrations of DC-HVR and EC-HVR taking it mixing ratios estimated by different methods are thus into account the mixing relationships between the consistent with one another, indicating that the esti- isotopic compositions of Sr, Nd and Pb and their mated elemental and isotopic compositions of DC- reciprocal concentrations (Table 2). The estimated HVR and EC-HVR are satisfactory. compositions (Table 2) indicate that EC-HVR has a granodioritic composition and DC-HVR has a basaltic 5.3.EC-HVR:an old subcontinental lithosphere? composition. This result is consistent with the Sr, Nd and Pb isotopic signatures of both components. The The Sr and Nd isotopic compositions of DC-HVR whole-rock composition of DC-HVR is classified as and HVR resemble those of the Yamato Basin and high-alumina or low-K basalt, according to their SiO2, Ulreung and Dog islands rocks (Fig. 5), suggesting Na2O and K2O contents (Kuno, 1966; Peccerillo and similar provenance materials for both southern Taylor, 1976). This major element signature of DC- Kyushu and the southern area of the Sea of Japan. HVR corresponds well to the Yamato Basin basalt In the Sr-Nd isotopic diagram (Fig. 5), the volcanic (Kimura et al., 1987), which has intrinsic composition rocks in northern Taiwan (Chung et al., 1995), the before contamination with an EM1-like source. Philippine Islands (Castillo, 1996) and the southern Yanagi et al. (1991) found a linear correlation among Ryukyu Island arc (Shinjo, 1998) plot around and/or between HVR and DC-HVR. On the Pb isotopic clase) chemistries in andesite to dacite of the active diagrams (Fig. 4), HVR have higher 206Pb/204Pb than Mt. Sakurajima volcano (Fig. 1), and proposed a volcanic rocks from northern Taiwan and Ulreung binary mixing between primary basaltic magma and and Dog islands. However, volcanic rocks in the residual dacitic magma for the genesis of Sakurajima Philippine Islands plot between HVR and the volcanic rocks. The major element compositions of Yamato Basin rock, and those of northern Taiwan the two magmas given by Yanagi et al. (1991) are similar to those of EC-HVR and DC-HVR (Table 2). (Fig. 4). All these volcanic rocks have higher Basalt from the Hisatsu volcanic province also has similar elemental compositions to those of DC-HVR the Dupal-like Pb isotopic signature. Tatsumoto and (Table 2). These results confirm that our binary Nakamura (1991) proposed that an old subcontinen- mixing model can be regionally and temporally ex- tal lithosphere with high Th/U and lithophile ele- tended to volcanic systems in southern Kyushu. ments, produced by a long geological time (>1 Ga), We can calculate the mixing ratio of EC-HVR or was responsible for the formation of Ulreung and DC-HVR for each HVR andesite using the isotopic Dog islands rock with EM1-like and Dupal-anomaly and element compositions in Table 2. It can be seen signatures. from the Sr-Nd isotopic diagram (Fig. 2) that the A recent seismic tomographic study (Sadeghi et contribution of EC-HVR in the Kurosonsan group al., 200o) has shown that strong low-velocity P- 32 T. Hosono et al. / Chemical Geology 201 (2003) 19-36 wave anomalies develop off western Kyushu at a 5.4.Evolution of HVR and constraint on mineraliza- depth of approximately 40-100 km. Although the tion-relatedmagmatism low-velocity anomaly corresponds to a low resis- tance zone (Shimoizumi et al., 1997) and is consid- It is likely from the above discussion that the lower ered to result from partial melting of an upwelling crust in southwestern Kyushu has granodioritic EC- hot mantle, the P-wave anomaly showing the HVR composition and was assimilated by hotter expected upwelling is not observed in the deeper basaltic magma with DC-HVR composition, leading mantle (Suzuki et al., 1999). It is pointed out from to the generation of the HVR (Fig. 7). The degree of these geophysical data that volcanic activity in assimilation of lower crust have increased with time western and southern Kyushu was not triggered by for the Kurosonsan group, while in the Shishimano dehydration of subducted oceanic crust as demon- group, it was less prominent in the SRD and SA-III. strated for volcanic rocks in northeastern Japan The SRD and SA-IlI showed higher crystallization (Nakada et al., 1997; Yanagi and Maeda, 1998; temperatures of phenocryst minerals of pyroxene and/ Seno, 1999). Interestingly, this low-velocity anomaly or hornblende than SA-I and SA-II (Hosono, 2003). at 40 km depth (Fig. 2 in Sadeghi et al., 2000) is This tendency indicates that the later-formed magma also observed in northern Taiwan, the southern in the Shishimano group contained larger amounts of Ryukyu arc, near Ulreung and Dog islands, and in high-temperature basaltic magma than the earlier- northwest Japan, where Ulreung and Dog islands- formed magma. The reverse zoning in pyroxene like or HVR-like volcanic rocks are distributed. In phenocrysts observed in SA-IlI would be caused by southern Kyushu the low P-wave anomaly is ob- the successive introduction of upwelling basaltic magma into more felsic magma. The presence of volcanic rocks are distributed (Fig. 3 in Sadeghi et reverse zoning of plagioclase phenocryst in SA-II al., 2000). We interpret from these studies that EC- may suggest that this hotter basaltic magma activity HVR-like materials correspond to the low velocity began from the SA-II stage. zone and are distributed regionally in the eastern Different from the HVR andesites, the two rhyo- margin of the Eurasian plate. dacites (KRD and SRD) deviate from the mixing According to a detailed seismic tomographic line connecting DC-HVR and EC-HVR (Fig. 6), study in southern Kyushu (Nishi et al.,2001),a indicating that they experienced fractional crystalli- low P-wave velocity zone develops at 5-25 and 60 zation. The larger deviation of SRD from the mixing km depth west of Mt. Kirishima. We deduce that line shows that the fractional crystallization was this low-velocity zone is almost correspondent to the more extensive in the SRD than in the KRD. subcontinental lithosphere in east Eurasia by Tatsu- Although SRD has the most felsic composition moto and Nakamura (1991). In southern Kyushu, it among HVR, its Sr-Nd-Pb isotopic signature probably constitutes the lower crust of granodioritic implies the dominant contribution of basaltic DC- composition, although we are uncertain as to wheth- HVR. This characteristic can be ascribed to a large er its age is older than 1 Ga as Tatsumoto and degree of fractional crystallization of anhydrous Nakamura (1991） considered because of a lack of mafic minerals (i.e. olivine and pyroxene) from a age data. This is consistent with the seismic study primitive andesite magma (Fig. 6). According to near Yakushima Island (Fig. 1) that the Ryukyu Hosono (2003), the crystallization temperature of Island arc, whose crustal structure resembles that phenocryst minerals of plagioclase and hornblende of southern Kyushu, comprises an upper-middle was higher about 100°C in SRD than in KRD. crust of the Shimanto group and its metamorphosed Despite the large degree of fractional crystallization, constituent of about 20 km thick, and a lower crust the large input of basaltic components led to the of about 10 km thick (Iwasaki et al., 1990). Iwasaki higher crystallization temperature of SRD compared et al. (1990) estimated that the lower crust near with that of KRD. Yakushima Island has a P-wave velocity of 6.7 km/ In the Kitami region of Hokkaido in northern s, suggesting that it may be composed of felsic Japan, basaltic to rhyolitic volcanic rocks of Miocene materials. age occur, and a total of 45 epithermal gold deposits T. Hosono et al. / Chemical Geology 201 (2003) 19-36 33 a a Late Miocene to Present Volcanic Rocks 2.4-0.5 Ma Middle Miocene Granitoids HVR KRD SRD (mineralization-related magma) SouthChinaSea PacificOcean 0 THHLATE ShimantoGroup Subcontinental Lower Crust Fractional Crystallization High-alumina fromPrimitive AndesiteMagma Basaltic Parent 50 Magma 100 Mantle Upwelling 150 Asthenosphere 200 300 100 0 (km) -a'in Fig.1. See text for details are hosted preferentially in the rhyolites. Kitami rhyolites are characteristically rare in hydrous miner- als (Watanabe et al., 1995). They have low 87Sr/86Sr Dupal isotopic anomaly, is distributed at the conti- nental margin of eastern Eurasia. We suggest that this compared with other volcanic rocks from northern lower crust has a granodioritic composition and may Japan (Takagi et al., 1999). In addition, dacitic rocks partly account for the low P-wave anomaly in this around the Chinkuashin gold deposit in northern region. A suite of HVR is generated by assimilation of Taiwan also show a characteristically unradiogenic lower crust with basaltic magma upwelling from the Sr isotopic composition similar to the SRD of HVR deeper mantle. Our model calculation shows that the (Chen, 1989). We consider that contribution of hot contribution of lower crust into the HVR increased basaltic magma and intense crystallization would have successively with age (from ca. 51% to 75%) in the been responsible for the formation of gold mineraliz- Kurosonsan group in the northern area, whereas those ing magma. in the Shishimano group in the southern area, the later-formed rhyodacite and andesite are smaller (ca. 57-58%) than the early-formed andesites (ca. 65- 6. Conclusions 68%). 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Nd and Sr isotopic Matsuura basalts in northwest Kyushu, Japan: an example of variations in acidic rocks formed under a peculiar tectonic envi- high-pressure open system fractional crystallization in a refilled ronment in Miocene Southwest Japan. Contrib. Mineral. Petrol. magma chamber near the crust-mantle boundary. Phys. Earth 99,1-10. Planet. Inter. 107, 203 -219. Tuttas, D., Habfast, K., 1982. High precision of lead isotope ratio Yanagi, T., Ichimaru, Y., Hirahara, S., 1991. Petrochemical evi- measurements. Finnigan Mat. Application Rep. 51. 19 pp. dence for coupled magma chambers beneath the Sakurajima Watanabe, Y., Kawano, Y., Takagi, T., 1995. Rhyolite related to the volcano, Kyusyu, Japan. Geochem. J. 25, 17-30. Miocene epithermal gold mineralization on the Kitami region, Zindler, A., Hart, S.R., 1986. Chemical geodynamics. Annual Re- Hokkaido Japan. Resource Geol. Spec. Issue 18, 237-248. view of Earth and Planetary Science 14, 493-571. Yanagi, T., Maeda, S., 1998. Magma evolution observed in the
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Journal of Asian Earth Sciences 188 (2020) 104107 Contents lists available at ScienceDirect Asia Journal of Asian Earth Sciences ELSEVIER journal homepage: www.elsevier.com/locate/jseaes Implications for boninitic magmatism in a late Cambrian nascent arc Toru Yamasaki Research Institute of Geology and Geoinforma tion,GeologicalSurvey of Japan,AIST,Tsukuba Central 7,1-1-1,Higashi,Tsukuba,Ibaraki305-8567,Japan ARTICLEINFO ABSTRACT Keywords: It is thought that the continental crust that ultimately became Japan (i.e., proto-Japan) formed during the early Ultramafic cumulate Cambrian near the margin of the South China block following its separation from the Australian block. However, Boninite the onset of timing of subduction and details of the magmatic evolution of the region are poorly understood, due Asaji ultramafic-mafic intrusion mainly to the sporadic ocurrence of Cambrian rocks. A high-temperature metamorphic complex is distributed Cambrian Clinopyroxene throughout eastern Kyushu island, Japan. At the northeastern end of the complex, a narrow zone of ultra- mafic-mafic rocks includes a late Cambrian ophiolitic assemblage, namely the Asaji ultramafic-mafic intrusion. Trace elements The ultramafic cumulates in the intrusion consist of harzburgite (spinel-olivine-orthopyroxene), Iherzolite-olivine websterite (olivine-orthopyroxene-clinopyroxene ± spinel), and websterite (orthopyroxene-clinopyroxene ± olivine), which formed in that order during differentiation of a parental magma. The chemical compositions of (Cr + Al)] ≤ 0.91 and TiO2 < 0.3 wt%, and clinopyroxene Ybn < 2) and comparable to those of spinel in bo- ninites. In addition, the calculated trace element composition of melts in equilibrium with clinopyroxene from the ultramafic rocks closely resembles the composition of the North Tongan boninite. These petrological and geo- chemical features suggest that the Asaji ultramafic-mafic intrusion consists of cumulates from a boninitic magma. The presence of boninitic magmatism suggests that the Asaji ultramafic-mafic intrusion was the result of sub- duction initiation and the infiltration of water into hot asthenosphere during the formation of nascent arc crust. 1. Introduction zircon gives a U-Pb age of 523-451Ma (Tsujimori et al., 2005; Kunugiza and Goto, 2010). The Oeyama Ophiolite in the Hida Marginal During the early Cambrian, the South China region, which had se- Belt includes harzburgite, although most is not particularly depleted parated from Australia, was located near proto-Japan (e.g., Isozaki (Cr# [atomic Cr/(Cr + Al)] = 0.4-0.6: Arai, 1980; Matsumoto et al., et al., 2010, 2015). The occurrence of metasomatic zircon in gneiss, and - d 1 granitoid magmatic rocks with c. 520-500 Ma ages, indicate that the fecture, which is also part of the Hida Marginal Belt, includes spinel arc-trench system of proto-Japan had developed by the middle Cam- grains with an average Cr# of 0.95 (Khedr and Arai, 2010). In the brian (Sakashima et al., 2003; Kunugiza and Goto, 2010; Tagiri et al., Kurosegawa Belt, peak ages of 600 Ma and 500-450 Ma have been re- 2010). However, the timing of the onset of subduction and the igneous ported for detrital zircons from metamorphic rocks in the Shikoku area processes are poorly understood due to the sporadic occurrence of c. (Yoshimoto et al., 2013), mafic granulites in the Kyushu area yield 500 Ma rocks. This uncertainty has prevented a precise reconstruction Sm-Nd isochron ages of 540, 490, and 420 Ma (Osanai et al., 2000), of the geodynamic setting for the formation of the proto-Japan arc. and metagabbroic rocks yield a zircon U-Pb age of 493-492 Ma (Osanai Lithologies key to understanding the geodynamic setting are peri- et al., 2014). In the Kurosegawa Belt, spinel with Cr# = 0.7-0.9 within dotites (now serpentinites), which represent fragments of mantle or s a s s lower crust. Nature of those ultramafic rocks strongly constrain the these rocks and the ultramafic rocks studied here has been noted by processes associated with the mantle-derived mafic magmas and those Soda and Takagi (2004). Finally, within the South Kitakami Belt, K-Ar geodynamic setting. Serpentinites from the area with an age of c. dating of hornblende in amphibolite yields an age of 500 Ma (Kanisawa 500 Ma are distributed throughout the Hida Marginal Belt (also referred et al., 1992), and U-Pb zircon dating of granitic rocks gives ages of to as Hida Gaien Belt), the Kurosegawa Belt (also referred to as the 500-490 Ma (Isawagawa tonalite) and 493 Ma (Shoboji diorite) (Isozaki Kurosegawa Tectonic Zone) and the South Kitakami Belt (Fig. la). A et al., 2015). Serpentinite bodies from the Hayachine-Miyamori area in Sm-Nd isochron from the Oeyama Ophiolite within the Hida Marginal Belt yields an age of 560 Ma (Hayasaka et al., 1995), and hydrothermal Yomogida, 1986; Machida and Ishiwatari, 2013). https://doi.org/10.1016/j.jseaes.2019.104107 Received 18 April 2019; Received in revised form 19 October 2019; Accepted 23 October 2019 Available online 31 October 2019 1367-9120/ @ 2019 Elsevier Ltd. All rights reserved. T.Yamasaki JourmalofAsianEarthSciences188(2020)104107 (a) N Honshu? 40°N South Sea of Japan Kitakami Belt Hida Marginal Suo Belt and Oeyama ophiolite Belt Akiyoshi Belt 35°N- Ryoke Belt Pacific Ocean Kyushu Shikoku Kurosegawa Belt (b) 200 km Fault 5 km *Syncline N ↑Anticline NTH01-04 NTH05 1km MtSHoji NTH06 Volcanic rocks Welded tuff Dike Rhyolite & graniteporphyry Onogawa Conglomerate & Group sandstone Watada granite Biotite granite Nioki granite Biotite granite Yamanaka Hornblende granodiorite granodiorite Quaternary Kinegaharu Quartz diorite& Sedimentary diorite diorite rocks Gabbro& Mafic Unit amphibolite Pyroclastic (Asaji ultra- Quartz- flow deposits mafic-mafic magnesite rock Neogene- intrusion) I Clinopyroxenite Quaternary Volcanic Serpentinite Hikata Unit Sandstone rocks Muddy melange Cretaceous coherent strata Ultramafic- mBasicvolcanic Permian accretionary complex mafic rocks Chokai Unit rocks Cretaceous granitic rocks Fault Chert W Cretaceous metamorphic rocks Fig. 1. Geological map of the study area. (a) Index map showing the location of the mapped area (b) (compiled from Ishiwatari and Tsujimori, 2003; Sakashima et al., 2003). (b) Geological map of the Oita-Taketa area showing the location of the mapped area (c), after Hoshizumi et al. (2015). (c) Geological map of the northern part of the Asaji Metamorphic Rocks (modified from Soda and Takagi, 2004). T.Yamasaki Jourmal of Asian Earth Sciences 188 (2020) 104107 A high-temperature metamorphic complex, termed the Asaji metamorphism due to the intrusion of Cretaceous plutonic rocks (Ryoke- Metamorphic Rocks, as the name of geologic body, is distributed from type metamorphism) (Miyazaki et al., 2014). Megusuno in Oita City to Ueki in Taketa City, Oita Prefecture, Kyushu, The ultramafic-mafic rocks in the Ultramafic Unit are distributed from Japan. The complex is around 4 km wide and extends for 25 km from NE Megusuno to Irikura, south of Oita City, in a 1-km-wide zone extending for to SW (Fig. 1a, b) (Ono, 1963; Oshima et al., 1971). A narrow zone of 6 km in a NE-SW direction. The rocks also occur as a small body (500 m ultramafic-mafic rocks is developed at the northeastern end of the com- wide, extending for 2km in a N-S direction) along the Otoge Fault in plex. This zone includes ophiolitic rocks such as serpentinized peridotites, Sawada, Ono Town, Bungo-ono City (Fig. 1c). In addition, a small body pyroxenites, gabbros, and plagiogranites. The plagiogranites yield a U-Pb occurs in Asaji Town, Taketa City (shown as metamorphic rocks in Fig. 1b) zircon age of 497 ± 3 Ma (Miyazaki et al., 2014; Hoshizumi et al., 2015). (Ono et al., 1977). The ultramafic-mafic rocks are composed mainly of The ultramafic-mafic rocks in the Asaji area contains chromian spinels peridotites and clinopyroxenites, with minor gabbros (Hayasaka et al., with Cr# ≤ 0.9 and TiO2 < 0.5 wt% (Soda and Takagi, 2004). As spinels 1989; Soda and Takagi, 2004). Among these lithologies, the peridotites are with such compositions are typically found within boninites (e.g., Kuroda completely serpentinized and exhibit a scaly texture (Hayasaka et al., 1989). et al., 1978; Walker and Cameron, 1983; Bloomer and Hawkins, 1987; Quartz-magnesite and tremolite rocks, which presumably formed by the Falloon et al., 1989; Arai, 1992; van der Laan et al., 1992; Sobolev and alteration and metamorphism of ultramafic rocks, also occur in this zone Danyushevsky, 1994; Ishikawa et al., 2002). (Hayasaka et al., 1989; Soda and Takagi, 2004; Fuji et al., 2008). According to Soda and Takagi (2004), spinel and clinopyroxene grains within the ultramafic-mafic rocks in the Asaji area are fresh and 3. Samples and analytical methods preserve their primary compositions, although the degree of serpenti- nization of the rocks is generally substantial. In particular, the trace The studied samples were collected from the areas of ^clinopyrox- element composition of clinopyroxene can preserve important in- enite' shown by Soda and Takagi (2004), from Megusuno and Irikura, formation regarding the geochemical characteristics of the parental south of Oita City (Fig. 1b). In addition to five samples from these areas, melts with which they equilibrated (e.g., Ross and Elthon, 1993; a serpentinite from Sawada, Oita City (southwest of Irikura; Fig. 2b), Kelemen et al., 1995). Here, I report the trace element compositions of was also examined. The locations of the samples are shown in Fig. 1b. clinopyroxene as well as the major element compositions of clinopyr- While rock names based on mineral assemblages, such as clinopyrox- oxene and spinel, and whole-rock compositions of ultramafic-mafic enite and serpentinite, have been used in earlier studies, the names used rocks in the Asaji area, and discuss the petrogenesis and geodynamic here follow Le Maitre (2002). setting of these rocks in the late Cambrian. For an estimation of the original mineral assemblages, Micro XRF (μ- XRF) mapping was carried out on thin sections to constrain elemental 2. Geological setting distributions. The mapping was conducted using a Bruker M4 TORNADO μ-XRF spectrometer in the Geological Survey of Japan Laboratory (GSJ- The pre-Upper Cretaceous basement rocks in the Asaji area, known as Lab), which uses an energy dispersive X-ray (EDX) detector to simulta- the Asaji Metamorphic Rocks, are classified into the three geological units, neously collect the energy spectra for all elements between Na and U from bottom to top: the Chokai Unit, the Ultramafic Unit, and the Hikata (Flude et al., 2017). The analytical conditions were as follows: Rh target, Unit (Fig. 1c) (Fuji et al., 2008). The Chokai Unit consists mainly of 40 kV accelerating voltage, 700 μA anode current, 25 μm beam diameter, metamorphosed pelitic, psammitic, and mafic rocks (greenstones), and 10 μm beam step size. Whole-rock major element compositions were whereas the Hikata Unit consists mainly of weakly metamorphosed sedi- measured with an X-ray fluorescence (XRF) spectrometer (PANalytical mentary rocks (Hayasaka et al., 1989; Teraoka et al., 1992; Fuji et al., Axios) in the GSJ-Lab using the method described by Yamasaki (2014). 2008; Miyazaki et al., 2014). The metamorphic rocks in the Chokai Unit The surfaces of chips from samples were scraped with a diamond disk to record contact metamorphism related to the emplacement of Cretaceous remove contamination from the rock saw. The chips were then cleaned plutonic rocks (Oshima et al., 1971; Karakida and Yamamoto, 1982; n n q Sasada, 1987; Itoyama and Takahashi, 1992; Osanai et al., 1993; Fuji dried in an oven for > 24 h. The dried samples were coarsely crushed in et al, 2008), which comprise diorite, granodiorite, granite, and gabbro u n s e pa u p r n s e (Hoshizumi et al., 2015). Solidification age of granodiorite and granite The quality of the XRF analyses was monitored through measurements of plutons ranges from 134.7 ± 2.8 Ma (Fuji et al., 2008) to 106 ± 4 Ma U.S. Geological Survey (USGS) geochemical reference material DNC-1. (Takagi et al., 2001, 2007). The pre-Upper Cretaceous basement rocks and The chemical compositions of the minerals were determined using a Cretaceous plutonic rocks are covered by pyroclastic flows and Quaternary JEOL JXA-8800R electron microprobe analyzer at the GSJ-Lab. The ac- conglomerate, and are in faulted contact with the Cretaceous Onogawa celerating voltage and beam current were 15 kV and 12 nA, respectively; Group (Matsumoto, 1936; Teraoka, 1970) in the northeast (Fig. 1b, c). ZAF corrections were applied to all analyses. Clinopyroxene trace ele- Interpretation of the geologic units that comprise the Asaji ment compositions were measured using laser ablation-inductively Metamorphic Rocks is complex. Originally, the Asaji Metamorphic Rocks coupled plasma-mass spectrometry (LA-ICP-MS) at the GSJ-Lab, which consists of a New Wave Research NWR213 laser ablation unit coupled to an Agilent 7700x quadrupole ICP-MS. Detailed information on the in- Zone of southwest Japan (e.g., Ono, 1963; Karakida and Yamamoto, s 1982; Osanai et al., 1993). However, following the discovery of Permian x o radiolaria in a metamudstone, it is evident that the Hikata Unit is a monitored using measurements of the NIST615 geochemical reference Permian accretionary complex and part of the Akiyoshi Belt (Toyohara material. A 100 μm laser spot size was used on sufficiently large clin- and Murata, 2003; Fuji et al., 2008; Fig. 1la). In addition, a U-Pb zircon opyroxene crystals, and 80 and 60 μm spots were used on smaller grains, age of 497 ± 3 Ma (late Cambrian) from plagiogranites in the Ultra- as appropriate. The laser energy fluence for 100, 80, and 60 μm was 6.5, mafic Unit has been reported from Asaji Town, Bungo-ono City 13, 18 J/cm2, respectively. (Miyazaki et al., 2014). Along with the Chokai Unit, both the Hikata Unit and Ultramafic Unit record weak high-T metamorphism during the 4. Petrography Cretaceous (Miyazaki et al., 2014). In addition, Fuji et al. (2008) and Hayasaka et al. (2013) found that the protolith of the Chokai Unit was 4.1. Lherzolite also part of the Permian accretionary complex. Thus, the Asaji Meta- morphic Rocks presumably consist of a late Cambrian ultramafic-mafic A sample of lherzolite (NTH-04) was collected from a roadside complex and a Permian accretionary complex that underwent high-T outcrop on the western flank of Mt. Shoji-dake, southwest of Megusuno, T.Yamasaki Jourmal of Asian Earth Sciences 188 (2020) 104107 a Fig. 2. Field occurrence of the ultramafic-mafic rocks in the study area. (a) Locality from which sample NTH-04 was collected. (b) Locality from which sample NTH- 06 was collected. (c) Locality from which sample NTH-03 was collected. (d) Locality from which sample NTH-07A was collected. Oita City (Fig. 1c). Although the highly weathered rocks appear to be and the rocks are fragmented into small pieces, typically a few cen- layered, details are obscured due to poor exposure (Fig. 2a). The rock is timeters across (Fig. 2b). The specimen is dark gray to gray in color, and dark in color and exhibits an aggregate texture comprising irregularly n q a p s shaped, yellowish brown grains. In thin section, the sample shows a crystals. In thin section, the sample consists of clinopyroxene grains of texture in which 90% consists of 0.5-1.0 mm rounded aggregates of talc 0.5-4.0 mm across, along with interstitial talc and magnetite (Fig. 3b, that are surrounded by small grains of magnetite, in which some grain c). The clinopyroxene is subhedral, shows a conspicuous cleavage, and boundaries are filled by clinopyroxene (Fig. 3a). Small amounts of eu- locally contains kink-bands. Oriented fine-grained magnetite, which hedral to subhedral chromian spinel (< 0.1 mm in diameter) also occur presumably defines an earlier cleavage, occurs in aggregates with fine- s'o) sudropnsd padeus-renueoy sapunoq uea sue grained talc and is interpreted as pseudomorphs after orthopyroxene long) are interpreted to have replaced orthopyroxene, most of which (Fig. 3b). Rounded mosaics of talc (1.0-1.5 mm across; Fig. 3c) and 0.5- show rounded shapes and are filled with fine-grained talc. This texture mm-sized rounded grains filled with serpentine are interpreted as closely resembles a former olivine cumulate composed mainly of olivine pseudomorphs after olivine. Smallquantities of fibrous tremolite occur crystals with minor orthopyroxene and clinopyroxene. The original within talc-rich portions of the rock. Magnetite occurs as anhedral, rock is interpreted to have been a lherzolite (i.e., an olivine-orthopyr- stubby prismatic grains (up to 0.5 mm long) within the talc-rich por- oxene-clinopyroxene cumulate). The Ca element map of this sample tion. The Ca maps shows substantial quantities (~33% for NTH-05 and reveals small amounts of clinopyroxene and its alteration products ~60% for NTH-06) of clinopyroxene and its alteration products (Fig. 4). surrounded by olivine and orthopyroxene (Fig. 4). The mesh-like dis- tribution of magnetite is clear in the Fe map (Fig. 4). Clinopyroxene 4.3. Olivine-bearing websterite grains occur as an interstitial aggregate, although individual grains are subhedral to euhedral cumulus minerals. Actinolite occurs as fine- A sample of olivine-bearing websterite (NTH-01) was also collected grained (< 0.5 mm) aggregates around clinopyroxene. from a roadside outcrop on the western flank of Mt. Shoji-dake, southwest of Megusuno, Oita City (Fig. 1c). The outcrop is heavily 4.2. Olivine websterite weathered, meaning that detailed lithological and structural observa- tions are difficult. On fresh surfaces the rocks are dark greenish gray Samples of olivine websterite (NTH-05 and NTH-06) were also and have a granular texture. In thin section, the rock comprises roughly collected from the roadside on the western flank of Mt. Shoji-dake, equal amounts of clinopyroxene crystals (~1.0 mm across), and southwest of Megusuno, Oita City (Fig. 1c). The outcrop is fractured rounded aggregates (~1.0 mm in size) of talc with abundant magnetite T.Yamasaki Jourmal of Asian Earth Sciences 188(2020) 104107 a NTH04 NTH-05 NTH-06 NTH-01 NTH-01 e NTH-07B NTH-03 6 NTH07A NTH07A Spl NTH07A Fig. 3. Photomicrographs of the ultramafic-mafic rocks from the study area. (a) Sample NTH-04; (b) sample NTH-05; (c) sample NTH-06; (d) & (e) sample NTH-01; (f) sample NTH-07B; (g) sample NTH-03; (h-j) sample NTH-07A. (a-d) and (g) plane-polarized light, (e-f) and (h-i) crossed-polarized light, (j) reflected light. FOV = 4.0 mm for (a-i), 2.0 mm for (j). Abbreviations: Cpx, clinopyroxene; Opx, orthopyroxene; Mt, magnetite; Ol, olivine; Tlc, talc; Tr, tremolite; Ser, serpentine; Hbl, hornblende; Pl, plagioclase; Spl, spinel. Parenthetic abbreviations denote original minerals of pseudomorphs. (granular to stubby prismatic in shape and < 0.5 mm in size), which are former cleavage (Fig. 3d, e). Elongate or lenticular clinopyroxene grains interpreted to be pseudomorphs after orthopyroxene (Fig. 3d). Within that probably represent exsolution lamellae are subparallel to this these altered grains, fine-grained magnetite is concentrated along a former cleavage (Fig. 3e). In addition to clinopyroxene and T.Yamasaki Journal of AsianEarthSciences188(2020)104107 NTH-04 NTH-05 NTH-06 NTH-01 NTH-07B NTH-07A Lherzolite Olivine Olivine Oivine-bearing Wehrlitic Harzburgite websterite websterite websterite Iherzolite Scan of thin section Ca Fe Fig. 4. Plane polarized light and X-ray maps of thin sections of the ultramafic-mafic rocks from the study area. The X-ray images show uncalibrated (semi- quantitative) elemental distributions based on total counts of Kα lines. Clinopyroxene and its alteration products (tremolite) appear as bright grains in the Ca map and dark (black) grains in the Fe map. Dark (black) grains in the Ca map are the alteration products of olivine and orthopyroxene, which are typically associated with tiny magnetite grains. orthopyroxene pseudomorphs, rounded grains (~0.5 mm in size) filled gabbroic rocks occur as lumps within a heavily weathered outcrop with greenish brown iddingsite probably represent pseudomorphs after (Fig. 2c), and might have been dikes or layers within the olivine web- sn s as s go d i sterite. The surfaces of fresh samples are dark greenish gray and have a clinopyroxene and its altered products, and dark portions comprising granular texture comprising green to dark gray amphibole and pyr- the alteration products of orthopyroxene and olivine (Fig. 4). These oxene, and pale-colored plagioclase. In thin section, the rock consists of darker portions in the Ca maps are relatively rich in Fe (Fig. 4). a large amount of amphibole with minor interstitial plagioclase and - n s a ss n o 4.4.Wehrlitic lherzolite dral (~0.5 mm long). Approximately 80% are green hornblende that is variably altered to chlorite. Other amphibole grains are brown horn- A sample of wehrlitic lherzolite (NTH-07B) was collected from the blende. In some cases, clinopyroxene is preserved in the cores of brown roadside in Sawada, southwest of Irikura-hikata, Oita City (Fig. 1c). hornblende (Fig. 3g). The plagioclase is subhedral to anhedral and fills Serpentinite is distributed throughout this outcrop, and the relationship interstices between amphibole. Whereas the rims of plagioclase are between wehrlite and serpentinite is unclear due to alteration, although generally fresh, the cores are commonly sericitized. Opaque minerals the wehrlitic lherzolite probably occurs as lenses or layers within the are anhedral and irregular in shape, and consist of Fe-Ti oxides. In serpentinite. The specimen is mostly black and contains small (< 1 some cases, aggregates of granular dark-gray clay minerals and lamellar mm) dark gray grains. The black portion is composed of serpentine, Fe-Ti oxides also occur. magnetite, bastite, and trace olivine. In thin section, the dark grey minerals are composed of aggregates of tremolite (Fig. 3f). The original 4.6.Harzburgite rock probably consisted of olivine, clinopyroxene, and minor ortho- pyroxene, and was probably a wehrlite, although the amount of original A sample of harzburgite (NTH-07A) was collected from the roadside orthopyroxene is uncertain due to the severe alteration. The original a few meters from sample NTH-07B (the wehrlite), close to the olivine and orthopyroxene cannot be identified from the X-ray maps. boundary between the‘serpentinite’ and‘clinopyroxenite’ zones of Soda These images show ~67% tremolite (Fig. 4), and there is a possibility and Takagi (2004). The rock is serpentinite with a dark bluish gray that the rock was a type of lherzolite. This is the only sample that color (Fig. 2d) and consists of serpentine and minor chromian spinel contains olivine, but fresh clinopyroxene is absent. The sample also (Fig. 3h). In thin section, olivine pseudomorphs (~0.1 mm in size) and contains trace amounts of small (< 0.1 mm) altered spinel grains. minor orthopyroxene pseudomorphs (bastite; ~0.5 mm long) are ob- served (Fig. 3j). Chromian spinel grains are mostly euhedral and com- 4.5.Pyroxene-hornblende gabbro monly have rims of magnetite (Fig. 3j). In general, chromian spinel in mantle peridotites shows an irregular amoeba-like irregular shape (or A sample of pyroxene-hornblende gabbro (NTH-03) was collected "holly-leaf shape; e.g., Le Mée et al., 2004), while those in the cumu- around 120 m from sample NTH-01 (the olivine websterite). The lates and gabbros occurs as euhedral and subhedral crystals (e.g., Arai T.Yamasaki Journal of Asian Earth Sciences 188 (2020) 104107 (continued on next page) 5 websterite roxene gabbrc 0 roxene 2 -hornblende gabbro 3 0 Repres i T. Yamasaki Jourmal of Asian Earth Sciences 188 (2020)104107 Iherzo ehrlitic 63 .n able1 (continued) 111 柳WWND T.Yamasaki Journal of Asian Earth Sciences 188 (2020) 104107 Table 2 were measured. The Cr# of spinel from samples NTH-04 and NTH-07A Representative microprobe analyses (wt%) of spinels in the ultramafic rocks is 0.98-0.89 and 0.91-0.79, and the Mg# is 0.10-0.02 and 0.60-0.14, from study area. respectively (Fig. 6a; Table 1). All spinel grains in NTH-04 and one spinel grain in NTH-07A have ferritchromite compositions (Fig. 6b). Lithology Lherzolite Harzburgite Sample# NTH04 NTH04 NTH07A NTH07A NTH07A under lower amphibolite to greenschist facies conditions, most of the other spinel compositions in sample NTH-07A show no evidence for Analytical# 101 106 5 6 20 SiO2 0.00 0.40 0.04 0.05 0.74 (Fig. 6b). With respect to Mg# and Cr#, all spinel grains other than TiO2 0.93 1.06 0.01 0.03 0.04 2.08 0.23 5.54 6.24 4.62 those with ferritchromite compositions have compositions similar to Al2O3 Cr2O3 25.23 21.90 59.79 58.55 56.27 those within boninites and forearc peridotites (Fig. 6c). In addition, Nio 0.08 0.01 0.01 0.04 0.05 spinel in sample NTH-07A is characterized by low TiO2 contents, si- FeO 27.21 27.67 21.80 22.47 22.55 milar to spinel grains in boninites (Fig. 6b; e.g., Arai, 1992). Fe2O3 40.60 46.78 3.87 4.31 7.51 MnO 2.67 2.36 1.04 1.02 1.18 MgO 0.46 0.30 6.18 5.86 5.27 Cao 0.00 0.01 0.01 0.00 0.02 5.2. Trace element composition of clinopyroxene Na2O 0.04 0.00 0.00 0.02 0.07 K20 0.00 0.00 0.00 0.01 0.02 The trace element compositions of clinopyroxene grains are listed in 99.30 100.72 98.29 98.60 98.33 Al 0.108 0.012 0.233 0.262 0.200 Table 3. Contents of V, Y, and Zr show a clear negative correlation Cr 0.874 0.777 1.688 1.650 1.632 ) # 1 Fe2+ 0.997 1.038 0.651 0.670 0.692 samples means that the high field strength elements (HFSEs) behaved as Fe3+ 1.339 1.579 0.104 0.116 0.207 incompatible elements within the crystallizing magma. However, the Mn 0.099 0.090 0.031 0.031 0.037 s n s no Mg 0.030 0.020 0.329 0.311 0.288 Ca 0.000 0.000 0.000 0.000 0.001 crystallized from a relatively evolved magma, whereas the other clin- Ni 0.003 0.000 0.000 0.001 0.002 opyroxene grains crystallized from more primitive magmas, as in- Cr/(Cr + Al) 0.890 0.984 0.879 0.863 0.891 s Mg/(Mg + Fe2+) 0.029 0.019 0.336 0.317 0.294 curved trend suggests changes in the partition coefficient for Sm be- FeO and FezOs were calculated assuming spinel stoichiometry. tween clinopyroxene and melt during magmatic differentiation. Primitive-mantle-normalized multi-element patterns and chondrite- normalized rare earth element (REE) patterns (normalizing values from et al., 2010). So-called ‘replacive dunite' has also been recognized in the Sun and McDonough, 1989) are shown in Fig. 8. For the multi-element mantle section of ophiolites, although this rock consists of olivine, patterns, samples NTH-01, NTH-04, NTH-05, and NTH-06 show similar spinel, and trace amounts of clinopyroxene with no orthopyroxene and/or parallel trace element patterns, although the highly in- (e.g., Kelemen et al., 1995). In addition, harzburgite tectonites usually compatible elements (Cs-U) show significant scatter. These samples contain small amounts of clinopyroxene, although there is no clin- commonly show negative anomalies in Ta, Pb, Zr, and Hf. Samples opyroxene or tremolite in this sample. The Ca maps show no Ca-rich NTH-04 and NTH-06 show a positive Sr anomaly, whereas samples minerals, except for within a lenticular xenolith-like portion which NTH-05 and NTH-01 do not. Since the total REE contents (EREE) in- presumably is a thin layer or fragment of a clinopyroxene-rich lithology crease in this order (i.e., NTH-04 < NTH-06 < NTH-05 < NTH-01), (Fig. 4). The compositional mapping sugests the sample is mainly this difference in Sr content presumably reflects the onset of plagioclase olivine and orthopyroxene, and is inferred to have been a spinel-oli- crystallization. The samples show a modest depletion in heavy REEs vine-orthopyroxene cumulate rather than a mantle tectonite. This in- (HREE) relative to light (LREE) and middle (MREE) REEs (e.g., the terpretation is not meant to imply that the serpentinites in the Ultra- Smn/Lun values of NTH-01, NTH-04, NTH-05, and NTH-06 are 3.45, mafic Unit are all of cumulate origin. 2.57, 3.33, and 2.71, respectively), except for the negative Zr and Hf anomalies in most samples. Sample NTH-03 shows flat middle to heavy 5. Mineral chemistry REE patterns, and both positive and negative Ti anomalies. 5.1.Major element composition of primary minerals s n n on n sd mus The major element compositions of the silicate minerals and spinel by low content of REEs on the whole. Sample NTH-03 shows relatively from the studied samples are listed in Tables 1 and 2, respectively. Fresh flat patterns and higher contents of Nd-Lu (~10 times chondrite va- olivine exists only in the wehrlite sample (NTH-O7B). The olivine yields lues), slopes down to the left for La-Nd, and a slight negative Eu Mg# (=100 × Mg/[Mg + Fe]) = 0.86-0.71, NiO = 0.14-0.01 wt%, anomaly (Fig. 8). and MnO = 1.52-0.26 wt% (Table 1). The Fo value in olivine was pos- sibly modified by contact metamorphism during intrusion of the Cre- taceous granitoids. The Nio content of olivine is lower, and the MnO 6. Whole-rock major element chemistry content is markedly higher, than in mantle olivine (e.g., Takahashi et al. 1987), which is consistent with lower primary Fo values relative to The major element compositions of the studied samples and re- mantle olivine. ference materials are given in Table 4. All samples except for NTH-03 Variations in the TiO2, MnO, NazO, and Cr2Os contents of clinopyr- show extremely low contents of TiO2, K2O, and P2Os, typical of ad- oxene against Mg# are shown in Fig. 5. Samples NTH-04 (Mg# = “(L8'0-68'0 = #8N) S0-HLN ‘(18'0-16'0 = #8N) 90-HLN ‘(68'0-76'0 centrations of AlzO3 in those samples reflect the absence of plagioclase. -1 A1 (040-08'0 = #8) S0-HN P ‘(18'0-8'0 = #8) 10-HN In addition, samples NTH-04, NTH-07A, and NTH-07B show relatively present a single magmatic trend with decreasing Mg# in the order given high MgO and low CaO contents, suggesting smaller quantities of (Fig. 5). However, plots of Cr2O3 content against Mg# apparently show clinopyroxene than in other samples. Sample NTH-03 (pyrox- two trends, one for #NTH-03 and one for the other samples. ene-hornblende gabbro) contains relatively high TiO2, Al2Os, Na2O, Due to the effects of alteration, spinel grains from only two samples K2O, and P2Os contents, suggesting it was enriched in melt. 9 T. Yamasaki Joumal of Asian Earth Sciences 188 (2020) 104107 口 NTH-01 (Ol-bg. websterite) 1.1 > NTH-03 (Px-Hbl gabbro) 1.0 TiO2 (wt%) Na2O (wt%) NTH-04 (Lherzolite) 0.9 NTH-05 (OI websterite) 0.8 0.4 XNTH-06 (OI websterite) 0.7 0.6 0.3 0.5 T 0.4 0.3 0.2 0.2 0.1 0.0 0.1 5.0- 26.0 Al2O3 (wt%) CaO (wt%) 4.0 24.0 B 3.0 22.0 2.0 20.0 1.0- 18.0 0.0 16.0 0.25- 0.8 MnO (wt%) Cr2O3(wt%) 0.20- 0.6 0.15 0.4 0.10- 0.2 0.05 X 0.00 品 0.0+ 0.65 0.70 0.75 0.80 0.85 0.90 0.95 0.65 0.70 0.75 0.80 0.85 0.90 0.95 Mg#[= Mg/(Mg +Fe)] Mg# [= Mg/(Mg + Fe)] Fig. 5. Plots of major oxides versus Mg# [= atomic Mg/(Mg + Fe)] for clinopyrox ne in the Asaji ultramaficmafic rocks. Ol, olivine; Px, pyroxene; Hbl, hornblende; Ol-bg., olivine-bearing. 7.Discussion suggest that all studied samples are cumulates. Thus, the igneous body consisting of those cumulates are collectively referred to hereafter as 7.1.Estimation of mineralmodes of primary phases the Asaji ultramaficmafic intrusion. In petrographic observations, the modal abundances of primary phases were estimated based on the re- Information on the petrographic observation and mineral chemistry lative abundance of clinopyroxene and pseudomorphs after olivine and 10 T.Yamasaki JournalofAsianEarthSciences188(2020)104107 Cr 1.0- (a) Omi serpentinites, (b) Hida Marginal Belt\ Happo-O'ne, Hida Marginal Belt 0.8 Boninites Omi serpentinites, Hida Marginal Belt r+Al) Forearc peridotites 0.6 Cr/(( /Happo-O'ne Lower Ⅱ / Dunite, Forearc amphibolite # Hida Marginal peridotites facies Belt 0.4 Hayachine & Miyamori, South Kitakami (484- Oeyama 421 Ma) Harzburgite, Greenschist Abyssal Hida Marginal Abyssal peridotites facies perido- Belt (560 Ma) 0.2 tites AI Fe3+ Hayachine & Miyamori, South Kitakami (484- 421 Ma) 3.5 10'0 (c) 1.0 0.8 0.6 0.4 0.2 0 3.0 Mg# = Mg/(Mg+Fe2+) 2.5 Intra-plate basalts 2.0 O NTH-04 (Lherzolite) 2 + NTH-07A (Harzburgite) O Hayachine & Miyamori, 1.5 South Kitakami (484- 一 Boninites 421 Ma) /MORB Asaji ultramafic-mafic 1.0 intrusion 一 (Soda & Takagi, 2004) 0.5 sland-arc basalts 0.0- 0.0 0.2 0.4 0.6 0.8 1.0 Cr# = Cr/(Cr+Al) Fig. 6. Compositions of spinel in the ultramafic-mafic rocks from the study area in terms of: (a) Cr# [= atomic Cr/(Cr + Al)] and Mg#; (b) Al-Cr-Fe3+ trivalent cation diagram; and (c) TiO2 and Cr#. The compositional fields for abyssal peridotites, forearc peridotites, and boninites in (a) are from Morishita et al. (2011) and references therein. The compositional fields of abyssal peridotites and forearc peridotites in (b) are from Khedr and Arai (2010) and references therein. The compositional fields of Omi serpentinites and Happo-O'ne and Oeyama in (a) and (b) are from Tsujimori (2004; Omi), Khedr and Arai (2010; Happo-O'ne), Arai (1980; Oeyama) and Kurokawa (1985; Oeyama). The compositional field of Hayachine & Miyamori in (a)-(c) are from Machida and Ishiwatari (2013). The curved gray arrow in (b) indicates the trend of changing spinel composition during retrogression from lower amphibolite to greenschist facies (Muintener et al., 2000). The field boundaries in (c) are from Arai (1992). orthopyroxene. While μ-XRF mapping can be a useful aid in estimating compositions of these minerals (Fig. 4). Therefore, the modes of the the modal abundance of clinopyroxenes, in this study olivine could not primary minerals were calculated using a multiple regression mass be distinguished from orthopyroxene because of the similar 11 T. Yamasaki Journal ofAsianEarthSciences188(2020)104107 websterite 03 (continued on NTH05(Olivine 000 00 3一 10 39一1 10 3 NTH04 (Lherzolite) 981 10 31 00 30 gabbro NTHO1 (Olivine-bearing 1 10 8一 8 (Pyrc 0.3 103 αのｍｈU ＧＯ T.Yamasaki JournalofAsianEarthSciences188(2020)104107 next page) 8 00 0 0 .0.0 0 (continued 0000 9. 4 0. 4 0. 12. 36 10. 3. 4 12 3. 7. 3 49. 19. 2. 4 2 0 8 82 2000 (oliv 0.72 10 9 NTH04 (Lherzolite) 86 00 8 2. 00 00 00 6 S (continued) Ｚ 昭盯ＺのＬ 13 T.Yamasaki Jourmal of Asian Earth Sciences 188 (2020) 104107 the ultramafic cumulates (i.e., the Generalized Petrological Mixing Model, GENMIX, of Le Maitre, 1979, 1981; recompiled by Kameya et al., 2001). 13849945159 For the mass balance calculations, constituent mineral chemistries were constrained as follows. For samples NTH-01, NTH-04, NTH-05, and NTH-06, the Mg# of the melt was calculated using an Fe:Mg par- tition coefficient (Kd) following Eq. (14) of Bédard (2010). The Mg# of olivine and orthopyroxene were then calculated using the Kd values of Beattie et al. (1991). Finally, the compositions of olivine and ortho- pyroxene were calculated stoichiometrically based on the Mg#. Since 800 o thopyroxene compositions were calculated from the highest Mg# of olivine in a similar way. In addition, as sample NTH-07A does not contain a primary silicate phase, the mineral compositions in sample NTH-07B (from the same outcrop) were used. The parameters for the mass balance calculation and estimated mineral modes of primary phases based on the mass balance calculation and μ-XRF mapping are provided in Table 5. 10 33 Based on the results of the mass balance calculation, the Asaji ul- tramafic-mafic intrusion can be subdivided into two groups (Fig. 9). The first group consists of relatively clinopyroxene-poor cumulates, namely harzburgite (NTH-07A), lherzolite (NTH-07B), and orthopyr- oxene-rich websterite (NTH-04). The other group consists of relatively clinopyroxene-rich olivine websterite (NTH-01 and NTH-06) and web- sterite (NTH-05). For samples NTH-04 and NTH-07B, the modes calculated from petrographic estimates are markedly different from those based on mass balance calculations (Table 5). In the case of sample NTH-04, the re- 2000 ance calculations is inconsistent with the petrographic observations. In vine this case, the modal abundances of orthopyroxene could have been (oliv overestimated in the mass balance calculation, as the calculated Mg# of 901 olivine was not very different to that in the other clinopyroxene-poor 一00 2 samples (i.e., NTH-07A and NTH-07B; Fig. 9). For sample NTH-07B, the modal abundance of clinopyroxene in the mass balance calculations was much smaller than that based on petrographic observations and μ- XRF mapping. In this case, the Cao content of tremolite (typi- cally ~ 13 wt%) is smaller than that of clinopyroxene (~23 wt%), which may have affected the estimated modal content of clinopyroxene. The large amount of tremolite identified in the μ-XRF mapping (~66.6%) cannot be fully explained. Regardless of this discrepancy in the estimates of modal abundance, the classification of this sample (i.e., that it is a type of lherzolite; see Section 4.4) is consistent with the mass balance calculations. The most 0145 important outcome of this estimate of mineral modes using multiple regression mass balance of the whole-rock and mineral compositions is the relationship between the mineral assemblages and the Mg# of the constituent minerals. The calculated Mg# of olivine suggests the for- mation of clinopyroxene-poor cumulates during an earlier stage of crystallization compared with the clinopyroxene-rich cumulates 10 10 716. 384. 25. 4 (Fig. 9). websterit 7.2.Nature of the parental melt and the order of crystallization in the ultramafic-mafic cumulates vine The compositions of chromian spinel from the Asaji ultra- mafic-mafic intrusion suggest the parental magma originated from a highly refractory mantle (Soda and Takagi, 2004; Fig. 6). In general, the Mg# of spinel increases with increasing temperature while the Cr# and Mg# of co-existing olivine remain constant (e.g., Fabries, 1979). Thus, the Mg# of spinel can constrain the petrogenesis of its host rocks. However, the Mg# of spinel studied here is highly variable. In parti- cular, spinel grains in lherzolite sample NTH-04 have extremely low Mg#. Spinel in sample NTH-04 is ferritchromite and does not preserve a primary composition. Soda and Takagi (2004) examined in detail the composition of spinel from the ultramafic rocks in the area, and found 14 T.Yamasaki Journal of AsianEarth Sciences188(2020)104107 1000 600 V(ppm) V (ppm) 500 400 100 300 200 100 10 0 100 25 Y (ppm) Y (ppm) 20 15 10 10 0 100 2.0 (wdd) IZ Sm (ppm) 1.5 口 1.0 0.5 0.0 0.70 0.75 0.80 0.85 0.90 0.95 0 5 10 15 20 25 Mg#[= Mg/(Mg +Fe)] Zr (ppm) Fig. 7. Trace element composition of clinopyroxene in the ultramafic-mafic rocks from the study area plotted against Mg# [= Mg/(Mg + Fe)] and Zr. The symbols are as in Fig. 5. that the Mg# decreased from core to rim due to serpentinization and forearc mantle, but similar to boninites with a similar Cr#. In addition, contact metamorphism. Therefore, lower Mg# values in spinel grains in the relationship between Cr# and TiO2 content indicates that the spinel samples NTH-04 and NTH-07A probably reflect modification. Con- originated from a boninite-like depleted magma. versely, the high Mg# of spinel in sample NTH-07A is likely a primary The contents of HFSE in clinopyroxene increase with decreasing Mg#, similar to REE (Figs. 6 and 7). These features suggest that the 15 T.Yamasaki JournalofAsianEarthSciences188(2020)104107 Primitive mantle-normalized Chondrite-normalized 100 100- NTH-01 (Ol-bg. websterite) NTH-01 (OlI-bg. websterite) ophiolitic slices overthrusted onto the blue- 0.01 100 100 NTH-03 (Px-Hbl gabbro) NTH03 (Px-Hbl gabbro) reversed zoning is observed; jadeite molecule 0.01 NTH-04(Lherzolite NTH-04 (Lherzolite) 10 0. 0.01 100 has been revealed in the Horokanai area. NTH-05 (Ol websterite) NTH-05 (Ol websterite) north region it locates close to the Kamui- 10 change of amphibole from core.to rim in its 0.01 100g 100 NTH-06(Ol websterite) NTH-06(O| websterite) blueschists. The root of the ophiolitic rocks is 0.01 Cs BaUTa CePr Nd Zr Eu Gd Dy HoTm Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu RbThNb LaPbSrSmHf TiTbYErYb Fig. 8. Primitive mantle- and chondrite-normalized trace element diagrams for clinopyroxene from the ultramafic-mafic rocks in the study area. The normalizing values are from Sun and McDonough (1989). Ol, olivine; Px, pyroxene; Hbl, hornblende; Ol-bg., olivine-bearing. contents of HFSE and REE in both clinopyroxene and magma increase adjacent to the western margin of the Ultramafic Unit decreases from with ongoing magmatic differentiation. When plotted on the isotherms the granite pluton to the eastern Ultramafic Unit (Asaji ultra- proposed by Lindsley and Andersen (1983), the composition of clin- mafic-mafic intrusion). The highest-temperature (biotite zone) condi- opyroxene grains is consistent with equilibration temperatures of tions near the contact with the pluton are estimated at 470-660 °C 500-1100°C (Fig. 10). The degree of metamorphism of the Chokai Unit (Fuji et al., 2008). Since the ultramafic-mafic rocks of this study are in T.Yamasaki Journal of AsianEarthSciences188(2020)104107 Table 4 Whole rock major element (wt%) compositions of the rocks from the Asaji ultramafic-mafic intrsuion and reference material. Lithology Ol-bg. Px-Hbl Lherzolite Olivine Olivine Harzburgite Wehrlitic Reference material Websterite gabbro websterite websterite lherzolite Sample# NTH-01 NTH-03 NTH-04 NTH-05 NTH-06 NTH-07A NTH-07B DNC-1 DNC-1 R.V. SiO2 50.32 49.70 54.91 54.65 52.87 45.50 43.68 47.58 47.20 TiO2 0.16 1.57 0.04 0.09 0.07 0.05 0.03 0.49 0.48 Al2O3 3.05 14.76 1.12 1.26 1.47 2.69 0.81 18.58 18.36 Fe2O3* 12.60 12.99 9.68 7.38 7.32 8.88 16.71 10.03 9.98 MnO 0.20 0.22 0.22 0.21 0.17 0.07 0.19 0.15 0.15 MgO 21.05 6.46 30.31 21.44 22.13 43.37 37.68 10.23 10.14 Cao 12.75 10.08 2.47 14.16 15.96 0.02 1.77 11.43 11.50 Na2O 0.20 4.07 0.09 0.23 0.20 0.07 0.06 1.96 1.89 K20 0.01 0.30 0.01 0.02 0.01 0.01 0.01 0.23 0.23 P205 0.01 0.13 n.d. 0.01 0.01 <0.01 <0.01 0.07 0.07 Total 100.35 100.27 98.86 100.03 100.52 100.19 100.92 100.74 100.00 LO1 2.86 1.58 4.82 1.03 2.27 12.39 7.93 Fe2O3* denotes total Fe as Fe2O3. DNC-1 R.V. is reference values from Gladney and Roelandts (1987). n.d., not determined; LO1, loss on ignition. Ol; olivine, Px; pyroxene, Hbl; hornblende. Table 5 Parameters for mass balance calculation and calculated mineral modes for ultramafic rocks from the study area. Sample# NTH-01 NTH-04 NTH-05 NTH-06 NTH-07A NTH-07B Cpx Mg# 0.83 0.90 0.88 0.89 0.91 0.91 Kd (eq. 14) 0.33 0.35 0.35 0.35 0l Mg# 0.73 0.85 0.81 0.82 0.86 0.86 Opx Mg# 0.75 0.85 0.82 0.83 0.87 0.87 Cpx mode (X-ray mapping, vol.%) 62.11 12.99 56.69 56.01 0.54 66.60 Cpx mode (mass balance, wt%) 58.72 14.53 61.86 69.23 0.00 8.49 Ol mode (mass balance, wt%) 15.71 4.95 0.00 9.87 59.64 58.28 Opx mode (mass balance, wt%) 25.57 80.29 38.14 20.90 39.38 27.54 Spl mode (mass balance, wt%) 0.22 0.98 5.68 Mg# = Mg/(Mg + Fe). Cpx Mg#s for NTH-01, NTH-04, NTH-05 and NTH-06 were averaged composition of clinopyroxene in each samples. Cpx Mg# for NTH-07B were calculated from Ol Fo, and Cpx Mg# for NTH-07A were extrapolated those in the NTH-07B. Kd (Eq. (14) is Fe = Mg exchange coefficient between clin- opyroxene and silicate melts calculated using Eq. (14) in Bédard (2010). Ol Mg#s and Opx Mg#s were calculated from Mg#s of melt in equilibrium with clin- opyroxene, using Fe = Mg exchange coefficient between olivine and silicate melts, and orthopyroxene and silicate melts (Beattie et al., 1991), respectively. contact with cooler muscovite zone rocks (Fuji et al., 2008), the esti- 10 mated equilibrium temperature based on the compositions of clin- opyroxene is unlikely to be associated with contact metamorphism. Thus, temperatures in excess of 750°C probably reflect magmatic crystallization temperatures. When comparing the measured REE compositions of clinopyroxene with those from mafic rocks with similar Mg# from the Hess Deep and East Pacific Rise fast-spreading ridges, which are highly depleted, the NTH-07A;- NTH-07B; studied samples show significantly lower concentrations of MREE to Fo86 Fo86 n lherzolite magmas that were markedly depleted in MREE-HREE, consistent with the compositions of chromian spinel. In addition, the TiO2 contents of olivinewebsterite clinopyroxene from the study area (in NTH-06, TiO2 < 0.15 wt%) are markedly lower than in clinopyroxene from the Hess Deep NTH-01;F073 (0.26-0.46 wt%; Co0gan et al., 2002). NTH-06;F082 10 To constrain the chemical composition of the magma, equilibrium melt compositions were calculated using the trace element chemistry of websterite the clinopyroxene in combination with appropriate mineral-melt par- Opx NTH-04; F085 Cpx tition coefficients as described below. The depleted nature of clin- NTH-05;F081 opyroxene and spinel, and the presence of hornblende inclusions within Fig. 9. Calculated mineral modes for ultramafic rocks from the study area. Fo denotes the calculated olivine forsterite content [= 100 × atomic Mg/ oxene in the studied samples coexisted with a hydrous parental magma. (Mg + Fe)]. The methods and parameters for the calculations are given in However, partition coefficients for REE between clinopyroxene and Table 5 and the text. The modal abundances were converted from wt% to vo- hydrous melt at low pressures are poorly understood. Here, experi- lume %. The dashed gray arrow shows the estimated order of crystallization of mentally-determined partition coefficients for melting of hydrous the cumulus minerals. Ol, olivine; Opx, orthopyroxene; Cpx, clinopyroxene. The mantle (1.3 GPa; McDade et al., 2003) are tentatively used for the nomenclature of ultramafic rocks follows Le Maitre (2002). 17 T.Yamasaki Jourmal of Asian Earth Sciences 188(2020) 104107 highest Mg# clinopyroxene grains, which occur in sample NTH-04 mineral compositions in the ultramafic rocks from the study area, accord (Mg# = 0.92-0.89). The difference in clinopyroxene-melt partition with the phenocryst assemblage observed in boninites. Thus, it can be coefficients between the lower crust and mantle is unlikely to sig- concluded that the parental magma for the Asaji ultramafic-mafic intrusion nificantly affect the calculations (e.g., Bédard, 2014). A detailed eva- had a boninitic composition. luation of the partition coefficients is discussed later. The calculated equilibrium melt composition shows weakly fractio- nated REE patterns (e.g., Lan/Ybn = 4.19 on average; Fig. 11b). Although 7.3. Geodynamic setting for formation of the ultramafic-mafic cumulates typical boninites shows characteristic concave-up REE patterns (e.g., Taylor et al., 1994; Fig. 11b), the north Tongan eastern group boninite U-Pb zircon dating of plagiogranites from the Asaji ultramafic-mafic (Falloon et al., 2007) shows LREE enrichment, similar to the calculated intrusion (Fuji et al., 2008) yields a late Cambrian age of 497 ± 3 Ma melt compositions (Fig. 11b). The north Tongan boninite is interpreted as (Miyazaki et al., 2014). Since the ultramafic rocks are lower-crustal cu- the product of rifting of the north Tonga Ridge oceanic arc (Sobolev and Danyushevsky, 1994; Falloon et al., 2008). Thus, if the partition coeffi- relationship between the cumulates and the plagiogranites is not cients are appropriate, the parental melt of the #NTH-04 clinopyroxene straightforward. However, the overall lithological assemblage is re- was derived from further HREE-depleted refractory mantle. garded as reflecting part of an ophiolite sequence (e.g., Ishiwatari and In general, an estimation of the equilibrium melt compositions using Tsujimori, 2003), in which the plagiogranites might be closely associated clinopyroxene-melt partition coefficients is problematic given the dif- with the ultramafic rocks (Miyazaki et al., 2014). Thus, it is possible that ferences between experimental and natural conditions. Qualitatively, the entire ophiolitic assemblage formed in the late Cambrian. clinopyroxene-melt partition coefficients for REEs under hydrous con- As mentioned earlier, serpentinites with an age of c. 500 Ma are ditions are lower than under anhydrous conditions (e.g., Gaetani et al., distributed throughout the Hida Marginal Belt, Kurosegawa Belt and 2003), and the partition coefficients become smaller with increasing South Kitakami Belt (Fig. la). Whereas the compositions of spinel pressure (e.g., Bédard, 2014). Taking account of the differences in grains in those serpentinites have not been significantly depleted (e.g., pressure and water content, the relatively low pressures and probable the spinels in the harzburgite from the Oeyama ophiolite; Arai, 1980; low activities of HzO during crystallization of the lower crustal cumulates Kurokawa, 1985: Fig. 6a), extremely depleted spinels have been re- would result in higher REE clinopyroxene-melt partition coefficients ported from several localities. Khedr and Arai (2010) examined the than the experimentally determined values. This suggests that the par- petrological features of the Happo-O'ne peridotite in the Hida Marginal ental magma of the ultramafic-mafic cumulates in the study area was Belt, and reported dunite with Cr# of 0.72 (Fig. 6a) that formed in a even more depleted than in the north Tongan and Izu-Ogasawara- forearc, in which hydration under relatively low temperatures during Mariana (IBM) forearc boninites, at least in terms of their HREE contents. the late-stage evolution resulted in the formation of tremolite-chlorite Based on the estimated mineral assemblages and Mg# of clinopyroxene, the ultramafic rocks in the study area formed in the following order: However, Machida and Ishiwatari (2013) suggested that at least some harzburgite [spinel-olivine-orthopyroxene cumulate] →>lherzolite-olivine of the small serpentinite bodies in the South Kitakami Belt were cu- websterite [olivine-orthopyroxene-clinopyroxene ± spinel cumulate] —→ websterite [orthopyroxene-clinopyroxene ± olivine cumulate]. This order indicated an origin within the sub-arc mantle, presumably in the of crystallization is supported by the estimation of the modes of the primary forearc. In addition to studies of peridotite, Nakama et al. (2010) ex- phases using multiple regression mass balance based on the whole-rock and amined U-Pb ages of detrital zircon from sandstones and river sands, mineral compositions of the ultramafic cumulates (Fig. 9). The inferred and sugested that the oldest age peak of 520-400 Ma, along with the order of formation of the cumulates implies a crystallization sequence of scarcity of Precambrian ages, indicated the existence of extensive spinel + olivine + orthopyroxene → olivine + orthopyroxene + clinopyr- granitic bodies that formed in an isolated intra-oceanic arc setting in oxene. Boninite is a high-MgO volcanic rocks that contains olivine and low- de-onod an pisns (ro) ' ia zos sdr-od Ca pyroxene phenocrysts. Plagioclase is restricted to the groundmass in n s evolved, slowly cooled examples (e.g., Crawford et al., 1989), and it has (south China) block. These discussions suggest that an arc-trench been assumed that the lack of plagioclase reflects depression of plagioclase system existed at around 500 Ma and that the refractory peridotites liquidus temperatures due to the presence of HzO. The estimated crystal- (serpentinites) were derived from the forearc. lization sequence of primitive magma, based on mineral assemblages and Under normal, steady-state subduction conditions, hydrous perido- tite in the forearc mantle wedge does not melt, and no magmatic rocks NTH-01 (Ol-bg. websterite) (oqqeb gaH-xd) E0-HIN NTH-04 (Lherzolite) NTH-05 (OI websterite) 50 △ <NTH-06 (Ol websterite) 500℃ -0.009 40 700℃ 900C 800°℃_ 1000°℃ >30 1100℃ 20 1200C 10 K ? K En 20 40 60 80 Fs sitions were calculated following the equation of Lindsley and Andersen (1983). Ol, olivine; Px, pyroxene; Hbl, hornblende; Ol-bg., olivine-bearing. 18 T.Yamasaki Jourmal of Asian Earth Sciences 188 (2020) 104107 100- (e.g., Umino and Kushiro, 1989; van der Laan et al., 1989). (a) E Boninitic magma is a distinctive magma formed by partial melting East Pacific Rise, of abnormally hot and hydrous mantle at low-pressure conditions. Hess Deep Cpx (Mg#89-82) Consequently, the existence of boninite is a reliable indicator of the tectonic setting (Crawford et al., 1989; Umino and Kushiro, 1989; van 10 der Laan et al., 1989; Bloomer et al., 1995). Boninitic magmas formed during repeated rifting in the Izu-Bonin-Mariana (IBM) arc and the Tonga arc (e.g., Falloon and Crawford, 1991; Sobolev and Danyushevsky, 1994; Danyushevsky et al., 1995; Ishizuka et al., 2006, 2011, 2014; Falloon et al., 2008; Reagan et al., 2008, 2010). In parti- curred at an early stage of magmatism after subduction initiation (e.g., Ishizuka et al., 2014). Given that the reported high-Cr# spinels are from NTH-06 Cpx (Mg#91-81) ~500 Ma bodies (Tsujimori, 2004; Khedr and Arai, 2010; Machida and Ishiwatari, 2013: Fig. 6) that are inferred to have formed in a forearc, it is suggested that boninitic magmatism in the study area is closely re- 0.1- lated forearc basalt (FAB) magmatism in the proto-Japan arc. Similar to the IBM, it can be assumed that the boninitic magmas in the study area La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu record the initiation of subduction and formation of the first oceanic arcs at c. 500 Ma. Kunugiza and Goto (2010) reported hydrothermal 100- (b) jadeitite including c. 520 Ma zircon grains from the Hida Marginal Belt. North Tongan eastern group boninite They assumed that the jadetite was formed at a newly developed oceanic arc and hot mantle wedge that received a high flux of fluids from the subducting slab, followed by oceanic spreading. This tectonic setting coincides with the assumed geodynamic setting for formation of 10- the ultramafic-mafic rocks in the study area. In addition, Osanai et al. (2014) reported high-pressure 493-492 Ma metagabbroic rocks from the Kurosegawa Belt in Kyushu and Shikoku. Based on the geochemical features of the metagabbroic rocks, Osanai et al. (2014) considered their protoliths to represent enriched mid-ocean-ridge basaltic magmas n temporal and spatial relationship to the assumed boninitic magmatism that is associated with the ultramafic-mafic rocks in the area. IBM forearc boninite magmatism are commonly attributed to the ascent of hot asthenosphere Calculated equilibrium melt with NTH-04 clinopyroxene from the deep mantle (e.g., Crawford et al., 1989). Such an environment 0.1- reflects migration of the spreading center associated with the ascent of La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu hot asthenosphere in a convergent setting, in which subduction supplies Fig. 11. Chondrite-normalized rare earth element patterns of: (a) clinopyroxene water to the mantle (Fig. 12) (e.g., Ishikawa et al., 2002; Yamasaki et al., from an ultramafic rock from the study area (sample NTH-06) and mid-ocean 2006). Since the lithological assemblage in the study area is essentially ridge gabbro; and (b) boninites and the calculated composition of melts in ophiolitic, the co-existence of MORB-type magmatic rocks (Osanai et al., equilibrium with clinopyroxene from the ultramafic rock in the study area. The 2014) supports an interpretation of subduction initiation at c. 500 Ma composition of clinopyroxene in a gabbro from Hess Deep, East Pacific Rise in (a) (Fig. 12). By contrast, there is a significant difference between the age of is after Coogan et al. (2002); the compositions of the North Tongan boninite and magmatism in the study area (c. 500 Ma) and that in the Hida Marginal Izu-Bonin-Mariana (IBM) boninite are from Falloon et al. (2007) and Taylor Belt (c. 520 Ma), although the geodynamic setting is assumed to be si- et al. (1994), respectively. The chondrite values are from Sun and McDonough milar, as discussed previously. If the uncertainties on the individual ages (1989). The equilibrium melt compositions in (b) were calculated using the are small, this age difference might suggest that the initiation of sub- partition coefficients of McDade et al. (2003); see text for detailed discussion. duction propagated from the proto-Hida Marginal Belt to the proto- Kyushu region, part of the present study area. are formed (e.g., Tatsumi and Maruyama, 1989). In anhydrous olivine tholeite magmas at relatively low pressures, olivine and plagioclase crystallize prior to clinopyroxene (e.g., Green and Ringwood, 1967; 8.Conclusions Grove et al., 1992). In basaltic systems, the cotectic relationship be- tween olivine and plagioclase breaks down at ≥ 5-7 kbar and olivine The composition of clinopyroxene and spinel grains from the Asaji and clinopyroxene crystallize prior to plagioclase (e.g., Kushiro and ultramafic-mafic intrusion indicates their derivation from a highly- Yoder, 1966; Green and Ringwood, 1967; Presnall et al., 1978). The depleted magma. The geochemical signature of clinopyroxene and spinel is similar to those in boninites. The calculated equilibrium melt Yorder and Tilley, 1962; Berndt et al., 2005). However, the sequences composition, using the trace element composition of clinopyroxene, arise consistent with a boninitic parental melt similar in composition to the North Tongan boninites. The order of crystallization of the Asaji hydrous, low pressure) cannot explain the crystallization of orthopyr- oxene followed by olivine, which requires partial melting of a highly magmas. Considering the previously reported age (c. 497 ± 3 Ma) for depleted mantle source or high-degree melting at low pressure (e.g.. the Asaji ultramafic-mafic intrusion, the geodynamic setting for the -u pue nou Aiaa ipun smooo Auo ues su uons (rooz osny formation of the parental boninitic magma is inferred to have been a drous conditions, which strongly depresses the liquidus temperature forearc rift, reflecting the early stages of magmatism following sub- duction initiation at around 520-500 Ma. 19 T.Yamasaki Jourmal of Asian Earth Sciences 188 (2020) 104107 (1) pre-arc stage (>500 Ma) Crawford, A.J., Falloon, T.J., Green, D.H., 1989. Classification, petrogenesis and tectonic setting of boninite In: Crawford, A.J. (Ed.), Boninite and related rocks. Unwin Hyman, London, pp. 1-49. E-MORB Coogan, L.A., Gillis, K.M., MacLeod, C.J., Thompson, G.M., Hekinian, R., 2002. 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Nappe structure of the Asaji metamorphic The authors declare that they have no known competing financial rocks, with special reference to geological structure of the basement complexes in Kyushu. The Memoirs of the Geological Society of Japan 33, 177-186 (in Japanese n e a n sn pd s with English abstract). ence the work reported in this paper. Hayasaka, Y., Sugimoto, T., Kano, T., 1995. Ophiolitic complex and metamorphic rocks in the Nimi-Katsuyama area, Okayama Prefecture. Excursion Guidebook of 102th Annual Meeting of the Geological Society of Japan, Hiroshima, 71-87 (in Japanese). Acknowledgements Hayasaka, Y, Kibayashi, K., Katsube, A., 2013. Terrane analysis for the basement complex of SW Japan and its tectonics: Implications of detrital zircon and monazite chron- The author is grateful to Kazuhiro Miyazaki (GSJ) for his comments on ological data. In: 12oth Annual Meeting of the Geological Society of Japan, Abstract, p. 80 (in Japanese). an earlier version of the manuscript. 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Mineral. flow deposit and their correlation to the Nioki Granite. J. Geol. Soc. Jpn 113, 1-14 (in Petrol. Sci. 42, 1-11. Japanese with English abstract). Matsumoto, I., Arai, S., Harada, T., 1995. Hydrous mineral inclusions in chromian spinel Takahashi, E.,Uto, K, Schiling, J.G., 1987. Primary magma compositions and Mg/Fe from the Yanomine ultramafic complex of the Sangun zone, Southwest Japan. J. ratios of their mantle residues along Mid-Atlantic Ridge 29N to 73N. Institute for Mineral., Petrol. Econ. Geol. 90, 333-338 (in Japanese with English abstract). Study of the Earth's Interior, Okayama University, Technical Report, Series A 9, 114. Matsumoto, I., Arai, S., Yamauchi, H., 1997. High-Al podiform chromitites in dunite- Tatsumi, Y., Maruyama, S., 1989. Boninites and high-Mg andesites: tectonics and petro- genesis. In: Crawford, A.J. (Ed.), Boninites. Unwin Hyman, London, pp. 50-71. Japan. J. Asian Earth Sci. 15, 295-302. Taylor, R.N., Nesbit, R.W., Vidal, P., Harmon, R.S., Auvray, B., Croudace, I.W., 1994. Matsumoto, T., 1936. Geology of the Onogawa Basin, Kyusyu. J. Geol. Soc. Jpn 43 Mineralogy, chemistry, and genesis of the Boninite Series Volcanics, Chichijima, 758-786, 815-853 (in Japanese with English abstract). Bonin Islands, Japan. J. Petrol. 35, 577-617. McDade, P., Blundy, J.D., Wood, B., 2003. Trace element partitioning between mantle Teraoka, Y., 1970. Cretaceous formations in the Onogawa basin and its vicinity, Kyushu, wedge peridotite and hydrous MgO-rich melt. Am. Mineral. 88, 1825-1831. southwest Japan. Geological Survey of Japan Report 237, 87 p (in Japanese with Miyazaki, K., Saito, M., Hoshizumi, H., Toshimitsu, S., Mizuno, K., Danhara, T., Iwano, H., English abstract). hirata, T., 2014. U-Pb and K-Ar ages of metamorphic and plutonic rocks in the area Teraoka, Y., Miyazaki, K., Hoshizumi, H., Yoshioka, T., Sakai, A., Ono, K., 1992. Geology of 1:200,000 quadrangle geological map of the Oita district. In: 121th Annual of the Inukai district. Geological Sheet Map at 1:50,000. Geological Survey of Japan, Meeting of the Geological Society of Japan, Abstract, p. 224 (in Japanese). 129 p (in Japanese with English abstract). Morishita, T., Dilek, Y., Shallo, M., Tamura, A., Arai, S., 2011. Insight into the mantle section of a maturing arc: the Eastern Mirdita ophiolite, Albania. Lithos 124, Annual Meeting of the Geological Society of Japan, Abstract, p. 80 (in Japanese). 215-226. Tsujimori, T., 2004. Origin of serpentinites in the Omi serpentinite melange (Hida Muintener, O., Hermann, J., Trommsdorff, V., 2000. Cooling history and exhumation of Mountains, Japan) deduced from zoned Cr-spinel. J. Geol. Soc. Jpn 110, 591-597 (in lower-crustal granulite and upper mantle (Malenco, eastern central Alps). J. Petrol. Japanese with English abstract). 41, 175-200. Tsujimori, T., Liou, J.D., Wooden, J., Miyamoto, T., 2005. U-Pb dating of large zircons in Nakama, T., Hirata, T., Otoh, S., Aoki, K, Yanai, S., Maruyama, S., 2010. Paleogeography low-temperature jadeitite from the Osayama serpentinite melange, Southwest Japan: of the Japanese Islands, Age spectra of detrital zircon and provenance history of the insights into the timing of serpentinization. Int. Geol. Rev. 47, 1048-1057. Orogen. J. Geogr. (Chigaku Zasshi) 119, 1161-1172 (in Japanese with English ab- Umino, S., Kushiro, I., 1989. Experimental studies on boninite petrogenesis. In: Crawford, stract). A.J. (Ed.), Boninites and related Rocks. Unwin Hyman, London, pp. 89-111. Ono, K., 1963. Kuju. Explanatory Text of The Gological Map of Japan, Scale 1:50,000, van der Laan, S.R., Flower, M.F.J., van Groos, A.F.K., 1989. Experimental evidence for the origin of boninites: Near-liquidus phase relations to 7.5 kbar. In: Crawford, A.J. (Ed.), Ono, K., Matsumoto, Y., Miyahisa, M., Teraoka, Y., Kambe, N., 1977. Geology of the Boninite and related Rocks. Unwin Hyman, London, pp. 112-147. Taketa District. Quadrangle Series, 1:50,000, Geological Survey of Japan, 156 p (in van der Laan, S.R., Arculus, R.J., Pearce, J.A., Murton, B.J., 1992. Petrography, mineral Japanese with English abstract). chemistry, and phase relations of the basement boninite series of Site 786, Izu-Bonin forearc. In: In: Fryer, P., Pearce, J.A. (Eds.), Proceedings of the Ocean Drilling morphic rocks, Oita Prefecture. Prof. Hisamichi Matsushita Memorial Volume, Program, Scientific Results, vol. 125. Ocean Drilling Program, College Station, Pp. 381-390 (in Japanese with English abstract). 171-201. Osanai, Y., Masao, S., Kagami, H., 1993. Rb-Sr whole rock isochron ages of granitic rocks Walker, D.A., Cameron, W.E., 1983. Boninite primary magmas: evidence from the Cape from the central Kyushu, Japan. The Memoirs of the Geological Society of Japan 42, Vogel Peninsula, PNG. Contrib. Miner. Petrol. 83, 150-158. 135-150 (in Japanese with English Abstract). Yamasaki, T., 2014. XRF major element analyses of silicate rocks using 1:10 dilution ratio Osanai, Y., Hamamoto, T., Kagami, H, Owada, M., Doyama, D., Ando, T.K., 2000. glass bead and a synthetically extended calibration curve method. Bull. Geol. Surv. Protolith and Sm-Nd geochronology of garnet-clinopyroxene granulite and garnet Japan 65, 97-103. amphibolite from the Kurosegawa Belt in Kyushu, southwest Japan. The Memoirs of Yamasaki, T., Maeda, J., Mizuta, T., 2006. Geochemical evidence in clinopyroxenes from the Geological Society of Japan 56, 199-212 (in Japanese with English Abstract). gabbroic sequence for two distinct magmatisms in the Oman ophiolite. Earth Planet. Osanai, Y., Yoshimoto, A., Nakano, N., Adachi, T., Kitano, I., Yonemura, K., Sasaki, J., Sci. Lett. 251, 52-65. Tsuchiya, N, Ishizuka, H, 2014. LA-ICP-MS zircon U-Pb geochronology of Paleozoic Yamasaki, T., Yamashita, K., Ogasawara, M., Saito, G., 2015. Multiple trace element granitic rocks and related igneous rocks from the Kurosegawa tectonic belt in Kyushu, analyses for silicate minerals and glasses by laser ablation-inductively coupled Southwest Japan. Japanese Mag. Mineral. Petrol. Sci. 43, 71-99 (in Japanese with plasma-mass spectrometry (LA-ICP-MS). Bull. Geol. Surv. Japan 66, 179-197. English abstract). Yorder, H.S.J., Tilley, C.E., 1962. Origin of basalt magmas: an experimental study of Presnall, D.C., Dixson, S.A., Dixson, J.R., O'Donnell, T.H., Brenner, N.L., Schrock, R.L., natural and synthetic rock systems. J. Petrol. 3, 342-532. Dycus, D.W., 1978. Liquidus phase relations on the join diopside-forsterite-anorthite Yoshimoto, A., Osanai, Y., Nakano, N., Adachi, T, Yonemura, K., Ishizuka, H., 2013. U-Pb from 1 atm to 20 kbar: their bearing on the generation and crystallization of basaltic detrital zircon dating of pelitic schists and quartzite from the Kurosegawa tectonic magma. Contrib. Miner. Petrol. 66, 203-220. Zone, Southwest Japan. J. Mineral. Petrol. Sci. 108, 178-183. 21
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Yamasaki (2020) - the 500 Ma Asaji ultramafic-mafic intrusion in Kyushu.txt
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Research Article
Blueschist-facies metamorphism during Paleozoic orogeny in
southwestern Japan: Phengite K–Ar ages of blueschist-facies tectonic
blocks in a serpentinite melange beneath early Paleozoic Oeyama
ophiolite
TATSUKI TSUJIMORI1AND TETSUMARU ITAYA2
1Department of Earth Sciences, Faculty of Sciences, Kanazawa University, Kanazawa 920–1192 and
2Research Institute of Natural Science, Okayama University of Science, Okayama 700–0005, Japan
Abstract Blueschist-bearing Osayama serpentinite melange develops beneath a peridotite
body of the Oeyama ophiolite which occupies the highest position structurally in the centralChugoku Mountains. The blueschist-facies tectonic blocks within the serpentinite melangeare divided into the lawsonite–pumpellyite grade, lower epidote grade and higher epidotegrade by the mineral assemblages of basic schists. The higher epidote-grade block is agarnet–glaucophane schist including eclogite-facies relic minerals and retrogressive law-sonite–pumpellyite-grade minerals. Gabbroic blocks derived from the Oeyama ophioliteare also enclosed as tectonic blocks in the serpentinite matrix and have experienced ablueschist metamorphism together with the other blueschist blocks. The mineralogic andparagenetic features of the Osayama blueschists are compatible with a hypothesis thatthey were derived from a coherent blueschist-facies metamorphic sequence, formed in asubduction zone with a low geothermal gradient ( ~10°C/km). Phengite K–Ar ages of 16
pelitic and one basic schists yield 289–327 Ma and concentrate around 320 Ma regardlessof protolith and metamorphic grade, suggesting quick exhumation of the schists at ca
320 Ma. These petrologic and geochronologic features suggest that the Osayamablueschists comprise a low-grade portion of the Carboniferous Renge metamorphic belt. The Osayama blueschists indicate that the ‘cold’ subduction type (Franciscan type)metamorphism to reach eclogite-facies and subsequent quick exhumation took place in the northwestern Pacific margin in Carboniferous time, like some other circum-Pacific oro-genic belts (western USA and eastern Australia), where such subduction metamorphismalready started as early as the Ordovician.
Key words: K–Ar phengite age, Osayama blueschist, Oeyama ophiolite, Paleozoic orogeny ,
Renge metamorphic belt, serpentinite melange.
Cotkin et al. 1992), eastern Australia ( ca 480 Ma,
Fukui et al. 1995) and Kurosegawa klippe, south-
western Japan ( ca 350–390 Ma, Ueda et al. 1980).
Paleozoic blueschists in these regions are alwaysassociated with Paleozoic ophiolite, and occur generally as tectonic blocks in a serpentinitemelange. Recent studies on Paleozoic ophiolites inthe circum-Pacific region have documented theophiolite formation in a supra-subduction zonesetting (forearc, volcanic arc, or back arc; Ozawa1988; Arai & Yurimoto 1994; Wallin & Metcalf1998). Tsujimori (1998) also showed that someINTRODUCTION
The blueschist-facies metamorphic rocks provide
critical evidence for paleo-subduction zones. In thecircum-Pacific orogenic belts, the incipient sub-duction of the paleo-Pacific plate took place duringthe Early–Middle Paleozoic, as indicated by theblueschist-facies metamorphic rocks from theKlamath Mountains, western USA ( ca 450 Ma,
Accepted for publication November 1998.
© 1999 Blackwell Science Asia Pty Ltd.The Island Arc (1999) 8,190–205
Paleozoic blueschist-facies metamorphism in SW Japan 191
ophiolitic fragments enclosed in the serpentinite
melange have experienced a blueschist metamor-phism together with other blueschist blocks. Thissuggests that the hanging wall of a subductionzone was also dragged into depths by tectonicerosion and metamorphosed.
In central Chugoku Mountains, southwestern
Japan, a serpentinite melange bearing blueschistblocks of various metamorphic grade (Osayamaserpentinite melange, Tsujimori 1998) developsbeneath the Early Paleozoic Oeyama ophiolite. This setting is a good example for studying subduction and exhumation processes through a joint geochronologic–petrologic method. Thispaper presents newly obtained K–Ar age data for
the blueschist-facies tectonic blocks from theOsayama serpentinite melange and discusses thetectonic implications of the Paleozoic orogeny insouthwestern Japan.
GEOLOGICAL SETTING
Southwestern Japan is a well-developed, circum-
Pacific type orogenic belt with oceanward growthof the accretionary complex since the middle Paleozoic (Fig. 1), as determined by a large amount of field-geological and biostratigraphic
Fig. 1 Geotectonic subdivision of Southwest Japan (modified from Isozaki & Itaya 1991; Isozaki & Maruyama 1991). Localities of the Pale ozoic
blueschists are also shown. Hd, Hida low-P/T metamorphic belt; Ok, Oki low-P/T metamorphic belt; Oe, Oeyama ophiolite; HM, Hida marginal belt; Rn,
Renge high-P/T metamorphic belt; Ak, Akiyoshi accretionary complex; Mz, Maizuru belt (Yakuno ophiolite); Ut, Ultra-Tamba accret ionary complex; So,
Suo high-P/T metamorphic belt; M-T, Mino-Tamba accretionary complex; Ry, Ryoke low-P/T metamorphic belt; Sb, Sambagawa high-P/T metamorphic
belt; Cn, northern Chichibu accretionary complex; Ks, Kurosegawa belt; Cs, southern Chichibu accretionary complex; Sh, Shimanto accretionary complex;
MTL, Median Tectonic Line; ISTL, Itoigawa–Shizuoka Tectonic Line.
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data (Hayasaka 1987; Nishimura 1990; Ishiwatari
1991; Isozaki & Itaya 1991; Isozaki 1996; Nakajima1997). Paleozoic ophiolite and blueschist have beensporadically distributed in the Chugoku Moun-tains, occupying the highest structural positions inthe nappe pile.
RENGE BLUESCHIST
Geochronologic data accumulated since the 1980sled to the subdivision of the ‘Sangun metamorphicbelt’ into two or three discrete units (Watanabe et al. 1987; Hayasaka 1987; Shibata & Nishimura
1989; Nishimura 1990; Isozaki & Maruyama 1991;Nakajima 1997). Most recently , Nishimura (1998)divided the ‘Sangun metamorphic belt’ into twobelts: the Renge belt (330–280 Ma) and the Suobelt (230–160 Ma). We follow the terminology ofNishimura (1998) for the high-P/T schist belts inthe Inner Zone of southwestern Japan. However,we distinguish the associated ophiolitic peridotitebodies from the Renge and Suo belts of Nishimura(1998) as ‘Oeyama ophiolite’, because they clearlypre-date the schists and have different tectono-metamorphic history .
The Renge blueschists in the Chugoku Moun-
tains occur as thin nappes, which are overlain bythe Oeyama ophiolite, and also appear as tectonicblocks within the serpentinite melange beneaththe Oeyama nappe (Fig. 2). The sporadic outcropsof the Renge blueschists comprise a disrupted
metamorphic belt, which has been considered asthe western extension of the blueschist-bearingOmi serpentinite melange. The Renge blueschistsmay have constituted a late Paleozoic regionalhigh-P/T metamorphic belt, which has been fragmented during exhumation and nappeemplacement.
OEYAMA OPHIOLITE
The peridotite bodies of the ‘Oeyama ophiolite’occupy the structurally highest position in theChugoku Mountains (Fig. 2). They are composedmainly of moderately depleted harzburgite (residual spinel peridotite) and dunite with gabbroic intrusions (diallage gabbro and dolerite).The eastern peridotite bodies such as at Oeyamahave slightly more fertile features than thewestern bodies, such as Tari-Misaka and Osayama(Arai 1980; Kurokawa 1985; Nozaka & Shibata1994; Matsumoto et al. 1995; Tsujimori 1998). The
podiform chromitites enclosed in dunite are char-acteristically developed only in western peridotite(Tari-Misaka body , Arai 1980; Matsumoto et al.
1997), and amphibolites (metacumulate and gneissose metagabbro) occur as tectonic block only in eastern peridotite bodies (Oeyama body , Kurokawa 1985; Wakasa body , Nishimura &Shibata 1989). The Ochiai–Hokubo body in the192 T . Tsujimori and T . Itaya
Fig. 2 (a) Distribution of geotectonic
nappe pile in eastern Chugoku Moun-tains. (b) Distribution of the peridotitebodies of the Oeyama ophiolite in thecentral Chugoku Mountains. Black areasrepresent ultramafic bodies.
14401738, 1999, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.1999.00231.x by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Paleozoic blueschist-facies metamorphism in SW Japan 193
central Chugoku Mountains, which is quite differ-
ent in petrologic features from the other bodies(Arai et al. 1988; Matsukage & Arai 1997), may be
a different geological unit in view of the radiomet-ric ages of gabbroic rocks (237–245 Ma, Nishimura& Shibata 1989).
The residual peridotites are mineralogically
very similar to the estimated mantle restites ofback-arc basin basalt and mid-ocean ridge basalt(MORB; Arai 1994). The presence of high-Al pod-iform chromitites and petrologic features of theresidual peridotite strongly indicate that theOeyama ophiolite represents a supra-subductionzone ophiolite formed in the back-arc basin orprimitive arc setting (Arai & Yurimoto 1994, 1995;Matsumoto et al. 1997; Zhou et al. 1998). The
gabbroic intrusions crosscutting peridotites of the Oeyama ophiolite show a MORB-like majorelement pattern (Hayasaka et al. 1995), which isalso compatible with the back-arc basin setting of
the Oeyama ophiolite.
GEOLOGY OF THE OSAYAMA SERPENTINITE
MELANGE
The Osayama serpentinite melange develops
beneath the Osayama peridotite body (Fig. 3). Theserpentinite melange is tectonically underlain by the Suo schists, and is in contact with theunmetamorphosed, molasse-type shallow marine sediments of the Jurassic Y amaoku Formation(Konishi 1954) on the north by a high-angle fault.All these rocks are unconformably overlain by theEarly Cretaceous Kyomiyama conglomerate. Themassive peridotite unit and the Suo schists haveundergone an overprint of contact metamorphismby Cretaceous granitic intrusives on the west.
Fig. 3 Geological map of the Osayama serpentinite melange (after Tsujimori 1998). Phengite K–Ar ages obtained in this present paper ar e also shown
in parentheses. Shaded area by broken lines represents the metamorphic zones in contact aureole by Cretaceous granites after No zaka and Shibata (1995).
LP , schist of lawsonite–pumpellyite grade; E, schist of epidote grade; Gb, diallage gabbro; Dl, dolerite; At, albitite.
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The blueschist-facies schists, fragments of the
Oeyama ophiolite (serpentinized peridotite,gabbro, dolerite) and metasomatic rocks (albitite,jadeitite, omphacitite, tremolite schist etc.) areenclosed as tectonic blocks of various size (10 cm to 1.5 km in length) in serpentinite matrixconsisting of schistose, friable and fine-grainedserpentinite with pebble-to-boulder-size frag-ments of serpentinized peridotite. The peridotiteblocks contain Fo
90.5–91.5 olivine, orthopyroxene with
2.4–3.0 wt% Al 2O3and chromian spinel with a
Cr/(Cr +Al) ratio of 0.40–57 as primary minerals.
Petrologic features of melange matrix peridotitesuggest that the melange matrix has been derivedfrom widely varying western peridotite bodies ofthe Oeyama ophiolite such as the Tari-Misaka andAshidachi bodies (Tsujimori 1998). Hashimoto andIgi (1970) first described lawsonite–glaucophaneschists in the eastern part of the serpentinitemelange here studied. The blueschist-faciesschists are divided into the lawsonite–pumpellyitegrade and epidote grades based on the mineralassemblages of basic schists intercalated in the pelitic schist. They correspond to the lawsonite–blueschist and epidote–blueschist faciesvarieties of Evans (1990), respectively . The epidotegrades contain two varieties, a garnet-free lower-grade block and a garnet-bearing higher-gradeblock (garnet–glaucophane schist). The blocks ofthe lawsonite–pumpellyite grade are the mostdominant type. The gabbro and dolerite blocksalso contain blueschist-facies mineral assemblagesof lawsonite–pumpellyite grade, but the gabbroicintrusives in the neighboring peridotite body donot have any blueschist-facies high-P/T minerals.The gabbroic blocks often grade into the basic schist of lawsonite–pumpellyite grade withincreasing textural deformation. The chemistry ofthe igneous clinopyroxenes and bulk rock compo-sitions of the Osayama gabbroic blocks indicatethat the blocks have been derived from the gabbroic intrusions of the Oeyama ophiolite (Tsujimori 1998).
PETROLOGY OF BLUESCHIST-FACIES BLOCK
LAWSONITE–PUMPELLYITE GRADE (LAWSONITE–
GLAUCOPHANE SCHIST AND GABBROIC FRAGMENTS)
The lawsonite–pumpellyite grade blocks are
characterized by the assemblage Na-amphibole +lawsonite or Na-amphibole +pumpellyite in
basic schists, although the mineral assemblage andtexture are variable from block to block. The basic
schists include the following mineral assemblageswith albite, quartz and titanite in excess: Na-amphibole +lawsonite +chlorite +phengite, Na-
amphibole +lawsonite +pumpellyite +chlorite,
Na-amphibole +lawsonite +pumpellyite +stilp-
nomelane, Na-amphibole +pumpellyite and
Na-amphibole +chlorite. In the fine-grained
sample, albite and quartz are exactly identified byusing an electron-probe microanalyzer. Titanite,relic augite, K-feldspar, sulfides, zircon and apatiteoccur as accessory minerals in some blocks. Most of the Na-amphiboles in this grade are glaucophane to ferro-glaucophane, and their compositions are variable for different mineralassemblages (Fig. 4). One block contains zonedNa-amphibole having a glaucophane core and a ferro-glaucophane rim (Fig. 4). Evidence of agreenschist-facies overprint such as an actinoliticrim on Na-amphibole is not observed in this grade.
The pelitic schists of the lawsonite–pumpellyite
grade consist mainly of quartz, phengite, chloriteand albite with minor titanite and apatite. Car-bonaceous matter, lawsonite, K-feldspar, tourma-line and carbonate minerals occur also in someblocks. A penetrative schistosity (S
1) defined by
phengite and chlorite is commonly observed. Insome cases, fine-scale crenulation cleavage (S
2) is
developed and overprints a crenulated S 1fabric.
Although phengite of S 1fabric is finer ( <0.2 mm)
than that of S 2(0.3–0.5 mm), no compositional dif-
ferences are recognized.
The ophiolitic fragments (gabbro and dolerite)
derived from the Oeyama ophiolite also have theblueschist-facies mineral assemblage, similar to lawsonite–pumpellyite-grade basic schists.Igneous plagioclase is replaced by aggregates ofpumpellyite or lawsonite and albite and igneousilmenite altered to aggregates of titanite. Na-amphibole occurs in three modes: overgrowingepitaxially on relict augite and hornblende, fillingcracks of clinopyroxene, and replacing patchedamphiboles included in clinopyroxene.
LOWER EPIDOTE GRADE (EPIDOTE–GLAUCOPHANE
SCHIST)
The constituent minerals of this grade are
commonly much coarser than the lawsonite–pumpellyite grade schists. The basic blocks arecharacterized by the assemblage Na-amphibole +
epidote +chlorite. The basic schist of the lower
epidote grade includes the following mineral assem-194 T . Tsujimori and T . Itaya
14401738, 1999, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.1999.00231.x by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Paleozoic blueschist-facies metamorphism in SW Japan 195
blages with albite, quartz and titanite: Na-amphi-
bole +epidote +chlorite, Na-amphibole +epidote +
chlorite +stilpnomelane, Na-amphibole +epidote +
pumpellyite and Na-amphibole +winchite +epidote
+chlorite +stilpnomelane. Na-amphiboles in this
grade are ferro-glaucophane to glaucophane, butsome blocks of the lower epidote grade contain prograde-zoned Na-amphibole having a crossitecore and glaucophane rim (Fig. 4). Albite oftenoccurs as porphyroblasts (maximum length: 3.5 mm). Although an actinolitic rim on Na-amphibole is rarely found in some small blocks,greenschist-facies overprinting is not observed inthis grade.
The pelitic schists of the grade contain mainly
chlorite, quartz, albite, and phengite with smallamounts of epidote and titanite. Albite commonlyoccurs as porphyroblasts (0.5–2.0 mm in length)which include tiny quartz, phengite, chlorite,apatite and rarely Na-amphibole. Na-amphibole,graphite, carbonate and garnet (Prp
1–2Alm 23–33
Sps 41–57Grs 19–25) are rarely observed. A penetrative
schistosity defined by coarse-grained phengite(0.5–0.8 mm in length) and chlorite is developed.
HIGHER EPIDOTE-GRADE BLOCK (GARNET–
GLAUCOPHANE SCHIST WITH ECLOGITIC MINERALASSEMBLAGE)
The higher epidote grade is defined by the coexist-
ence of almandine-rich garnet +glaucophane andthe presence of an eclogite-facies mineral assem-
blage. In the garnet–glaucophane schist, two dis-tinct blueschist-facies stages can be defined basedon the texture and mineral zoning. The peak meta-morphic stage is characterized by the assemblageNa-amphibole (glaucophane core) +garnet +rutile
+epidote +quartz +K-feldspar. The epidote por-
phyroblasts (maximum length: 2 mm) sometimesinclude eclogite-facies mineral assemblage, garnet+omphacite (Jd
35–50Di52–56Ae <9) +rutile +quartz +
glaucophane, as tiny inclusions ( <0.03 mm). The
retrograde stage is characterized by the assem-blage Na-amphibole (ferro-glaucophane rim) +
chlorite +pumpellyite +titanite ± phengite, which
is equivalent to the lawsonite–pumpellyite grade.In some cases, the strongly sheared phengite-richpart is developed in the outcrop. Although compo-sitional zoning from glaucophane core to ferro-glaucophane rim is common, such zoning is notobserved in the phengite-rich part (Fig. 4). Retro-grade ferro-glaucophane often fills cracks ofgarnet. Garnet porphyroblasts (up to 3 mm indiameter) often contain tiny inclusions of rutileand quartz. Garnets in the glaucophane-rich parthave higher (Mg +Fe) and lower Mn contents than
garnets in the garnet-rich layer (Fig. 5). Thegarnets show prograde zoning where Fe and Mg increase and Mn decreases from core to rim, and compositions of garnets within epidote por-phyroblast corresponds to the rim of those in theglaucophane-rich part (Fig. 5). The distributionFig. 4 Compositional variations of Na-amphiboles from the Osayama blueschists in Miyashiro’s diagram of Fe3+/(Fe3++Al) vsFe2+/(Fe2++Mg). The
arrows show compositional zoning (R, rim; C, core).
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coefficients of Fe and Mg, K DGrt-Cpx, between garnet
and omphacite in the epidote vary from 8.3 to 15.9.This garnet–glaucophane schist is a high-gradeblock which was overprinted with the other low-grade blueschists. More detailed petrology ofthe garnet–glaucophane schist will be describedelsewhere.
BULK ROCK COMPOSITIONS OF THE OSAYAMA PELITIC
SCHISTS
The bulk rock composition of a typical law-
sonite–pumpellyite-grade pelitic schist and threelower epidote-grade schists were analyzed. Noremarkable difference between the two grades was recognized. As compared with the average ofthe Sambagawa pelitic schists (Goto et al. 1996),
the Osayama pelitic schists are characterized byhigher MgO (2.9–3.8 wt%), FeO* (5.2–8.1 wt%),P
2O5(0.16–0.40 wt%), K 2O (3.3–5.6 wt%) and
moderate CaO (0.8–1.4 wt%), MnO (0.08–0.17 wt%)and Al
2O3(14.7–19.3 wt%). The MgO/(MgO +FeO*)
mole ratio is 0.39–0.45. The A ¢value of AFM
diagram [(Al 2O3–3K 2O-Na 2O)/(Al 2O3–3K 2O-Na 2O+
FeO* +MgO)] varies from –0.15 to 0.00 and issignificantly lower than that of the Sambagawa
average (0.11). The Osayama pelitic schists arericher in mafic components than the Sambagawapelitic schists.
METAMORPHIC CONDITIONS
The Na-amphiboles in the Osayama blueschists
are characterized by a low Fe
3+/(Fe3++Al) ratio,
and are in the glaucophane and ferro-glaucophanefields except for those in some lawsonite–pumpellyite-grade blueschist, and the core compo-sition of some zoned Na-amphiboles in the lowerepidote grade (Fig. 4). Phengites in the Osayamablueschists have Si contents significantly higherthan that in the underlying Suo pelitic schists (Fig. 6). The compositions of Na-amphibole andphengites of the Osayama blueschist showcommon high-P/T features.
Although any geothermometers based on Fe–
Mg exchange reactions are not applicable for the lawsonite–pumpellyite grade of the Osayamablueschists, its approximate P–T condition can bededuced by the mineral assemblage. In the lawsonite–pumpellyite grade, glaucophane +law-
sonite and glaucophane +pumpellyite assem-
blages are observed and albite is stable. TheSchreinemakers’ net for the NCMASH (Na
2O-
CaO-MgO-Al 2O3-SiO 2-H 2O) system shows that the196 T . Tsujimori and T . Itaya
Fig. 5 Composition of garnets of garnet–glaucophane schist of the
higher epidote grade in the Mn–Fe–Mg ternary diagram. Matrix garnet:(d), glaucophane rich; ( s), garnet-rich. Eclogite garnet: ( w), inclusions
within epidote.
Fig. 6 Compositional variations in Al vsSi (p.f.u. for O =22) for phen-
gites from the Renge blueschists and the Suo schists. (a) Pelitic schist;(b) basic schist. Osayama blueschists: ( d), lawsonite–pumpellyite grade;
(u), lower epidote grade; ( s), higher epidote grade. Suo schist: ( +),
pelitic schists (underlying the Osayama melange).
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Paleozoic blueschist-facies metamorphism in SW Japan 197
solid solution, gives a minimum pressure of 1.5 GPa
at ~550°C for omphacite (X Jd=0.35–0.50) +quartz
assemblage without albite. The P–T condition ofthe final retrogression stage may correspond tothat of the lawsonite–pumpellyite grade.
In the typical high-P/T type metamorphic belts,
such as the Franciscan, Kamuikotan and NewCaledonian, their low-grade portions are charac-terized by the common assemblage of glaucophane+lawsonite or pumpellyite, and then the glauco-
phane +epidote assemblage becomes gradually
stable with increasing metamorphic grade(Y okoyama et al. 1986; Maruyama & Liou 1988;
Takayama 1988; Shibakusa 1989). In the Osayamablueschist blocks, the paragenetic feature of thelawsonite–pumpellyite grade and lower epidote-grade blueschists may be interpreted as a coher-ent metamorphic sequence that has undergonetypical high-P/T type metamorphism in the sub-duction zone, with a geothermal gradient close to10°C/km (Miyashiro 1994). The presence of albiteindicates that the metamorphic condition liesbelow the jadeite–quartz reaction line. Originalcoherency of the metamorphic sequence for theOsayama blueschists is also supported from thegeochronologic data described in the followingsection.
K–AR AGE DETERMINATION
The K–Ar ages were determined for 20 phengite
separates from 16 metamorphic rocks (Table 1):two basic and pelitic schists from the lawsonite–pumpellyite grade, 12 albite porphyroblast-bearing pelitic schists from the lower epidotegrade, and two phengite-rich parts of a garnet–glaucophane schist (higher epidote grade).Mineral assemblages of the rocks dated are shownin Table 1.
Rock samples were crushed with a jaw crusher
and then sieved to obtain a proper grain-size forconcentrating phengite. The sieved fraction waswashed using de-ionized water and dried in anoven at 80°C. Phengites were concentrated usingan isodynamic separator and a tapping on a paper,and the collected phengite was treated with 2 mol/L HCl to dissolve out chlorite along cleavageplanes. The acid-treated sample was then washedrepeatedly with ion-exchanged water and dried at80°C.
The K–Ar age determination was carried out at
Okayama University of Science following Nagao et al. (1984) and Itaya et al. (1991). Potassium was
Fig. 7 K content (wt%) vsK–Ar age (Ma) diagram showing the effect
of grain size and impurities in phengites. Tie-lined data are from the samesample and the shaded marks represent the coarser grained phengite sep-arate. ( s), (150/200) lawsonite–pumpellyite grade; ( u) (150/200), ( )
(100/150), lower epidote grade; ( u) (150/200), ( ) (100/150), higher
epidote grade.
glaucophane +lawsonite assemblage is stable at a
higher pressure than the pumpellyite–actinolitefacies and pumpellyite–diopside facies (Banno1998). The glaucophane +pumpellyite stability field
also appears as a subfacies in the glaucophane +
lawsonite field in NCMASH systems (Frey et al.
1991; El-Shazly 1994). The P–T condition of the law-sonite–pumpellyite grade is restricted in a field ofthe lawsonite–blueschist facies where the glauco-phane +pumpellyite +albite assemblage is stable.
In the same sense, the P–T condition of the lowerepidote grade is limited to the albite-stable field in the epidote–blueschist facies. The petrogeneticgrid proposed by Evans (1990) indicates that the lawsonite–pumpellyite grade did not reach theclosure temperature ( ~350°C) of the K–Ar phengite
system, whereas those of the lower epidote-graderocks were probably above that temperature.
In the garnet–glaucophane schist of the higher
epidote grade, assuming that the inclusions withinepidote are in equilibrium, the garnet–clinopyrox-ene Fe–Mg exchange geothermometer by Krogh(1988) gives 530–620°C at 1.3 GPa. The geobarom-eter using the breakdown of low albite to jadeite +
quartz (Ghent et al. 1987), assuming ideal Jd–Di
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198 T . Tsujimori and T . Itaya
Table 1 Mineral assemblages of the samples used for the K–Ar age determination
Meta. Grade (Facies) LP (LBS) E (lower EBS) E (upper EBS)
Lithology Basic Pelitic Albite porphyroblast-bearing pelitic schist Grt–Gln schist
Sample OS162a OS80 OS182 OS277 OS304 OS190 OS224 OS329 OS93 OS350 OS188 OS318 OS267 OS281 OS23 OS23B
Na-amphibole • — • • —————————— • •
Lawsonite • — ———————————— — —
Epidote — — •••••••••••• • •
Chlorite • • •••••••••••• • •
G a r n e t —— —— • ————————— • •Phengite • • •••••••••••• • •
Quartz • • •••••••••••• • •
Albite • • •••••••••••• • •
Titanite • • •••••••••••• • •
Rutile — — ———————————— • •
Carbonaceous matter — • — — • ———— • — • — • — —
Others Kfs Kfs
K–Ar age (Ma)
[100/150] — — 324 — — — 318 319 292 — — — — — 322 319[150/200] 315 311 308 327 324 283 273 312 300 285 315 312 315 289
LP, lawsonite–pumpellyite grade; E, epidote grade; LBS, lawsonite–blueschist facies; EBS, epidote–blueschist facies.
Phengite K–Ar ages are also represented.
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Paleozoic blueschist-facies metamorphism in SW Japan 199
analyzed by flame photometry using a 2000 ppm
Cs buffer. Argon was analyzed on a 15-cm radiussector-type mass spectrometer (HIRU) having asingle collector system with an isotopic dilutionmethod using a
38Ar spike (Itaya et al. 1991). Decay
constants for 40K to 40Ar and 40Ca, and the 40K
abundance used in age calculation are 0.581 ·
10–10/year, 4.962 ·10–10/year, and 0.0001167, respec-
tively (Steiger & Jager 1977). The results are pre-sented in Table 2 and are also shown visually inFigs 7 and 8.
Phengite K–Ar ages of the Osayama blueschist-
facies tectonic blocks of the lawsonite–pumpellyite,lower epidote and higher epidote grade yield311–315, 273–327 and 289–322 Ma, respectively ,which are concentrated around 320 Ma as a whole.The phengite separates dated have potassium con-tents ranging from 5.52 to 8.89, and most of them(16 samples) are greater than 6.5 wt% in potash. Asmentioned earlier, the samples of the lawsonite–pumpellyite grade never reached the closure tem-perature ( ~350°C) of the K–Ar phengite system,
whereas those of the epidote-grade rocks wereprobably above that temperature. Itaya and Taka-sugi (1988) argued that the K–Ar phengite age ofthe low-grade Sambagawa schists, which have not experienced a culmination temperature higherthan the closure temperature of the phengite K–Arsystem, represented the timing of exhumation/cooling ages because of the argon depletion from
phengite by ductile deformation during theexhumation of the host schists. Thus, we interpretthe K–Ar phengite ages as the exhumation/coolingage soon after the blueschist-facies metamorphism.However, the range of K–Ar ages of the phengiteseparates from the blueschist-facies schist tectonicblocks, wider than analytical error of individualanalysis. They may be due to either or both of (i) thedifferent cooling age among the schists at the timebecause of different argon depletion processesduring ductile deformation in exhumation ofschists; and (ii) the effect of impurities in the phen-gite separates because some finer grained fractionshave significantly lower potassium content andyounger age (Fig. 7). Although the phengite K–Arages of the Osayama blueschists-facies schists havesome variation, the concentration at 320 Ma indi-cates that the exhumation of schists took placeapproximately at that time.
DISCUSSION
GEOLOGICAL SIGNIFICANCE OF THE OSAYAMA
BLUESCHISTS
In southwestern Japan, late Paleozoic high-P/T
schists are sporadically distributed (Fig. 1).Table 2 Phengite K–Ar age data of the blueschist-facies tectonic blocks from the Osayama serpentinite melange
Sample no. Fraction Potassium (wt%) Rad. argon 40 (10-8cc STP/g) Age (Ma) Air cont. (%)
Lawsonite–pumpellyite grade
OS162a 150/200 6.811 –0.136 9086 –88 314.7 –6.4 1.4
OS80 150/200 6.442 –0.129 8489 –82 311.0 –6.3 1.1
Lower epidote grade (albite porphyroblast-bearing pelitic schist)
OS182 100/150 8.204 –0.164 11311 –110 324.3 –6.6 0.4
OS277 150/200 5.517 –0.110 7184 –72 307.7 –6.3 0.8
OS304 150/200 8.211 –0.164 11418 –113 326.9 –6.7 0.9
OS190 100/150 8.435 –0.169 11602 –110 323.6 –6.6 0.5
OS224 100/150 7.470 –0.149 10068 –99 317.6 –6.5 0.5
OS329 100/150 7.819 –0.156 10582 –103 318.8 –6.5 0.7
150/200 7.422 –0.148 8810 –84 282.6 –5.8 0.9
OS93 100/150 6.946 –0.139 8542 –82 292.0 –6.0 0.4
150/200 6.184 –0.124 7083 –69 273.4 –5.6 0.7
OS350 150/200 7.942 –0.159 10480 –100 311.5 –6.3 0.6
OS188 150/200 8.891 –0.178 11254 –108 299.8 –6.1 0.7
OS318 150/200 7.657 –0.153 9168 –89 284.9 –5.9 1.1
OS267 150/200 7.550 –0.151 10084 –101 315.0 –6.5 0.7
OS281 150/200 6.684 –0.314 8840 –87 312.2 –6.4 0.8
Higher epidote grade (garnet–glaucophane schist)
OS23 100/150 7.483 –0.150 10240 –99 322.1 –6.5 0.7
150/200 6.721 –0.134 8960 –89 314.5 –6.4 0.8
OS23B 100/150 7.065 –0.141 9578 –92 319.4 –6.5 0.7
150/200 5.682 –0.114 6900 –66 288.6 –5.9 0.9
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Although they generally occur as tectonic blocks
in a serpentinite melange, a series of high-P/Tschists have been considered to be the constituentsof a late Paleozoic regional high-P/T metamorphicbelt, called the Renge belt (Nishimura 1998), basedon the geochronologic data shown in Fig. 8. Thephengite K–Ar ages from the Osayama blueschistsare within the age variation of the high-P/T schistsin the Omi, Wakasa, Toyogadake, Wakamiya andKiyama areas of the Renge belt, indicating that theOsayama blueschists is a constituent of the Rengebelt.
Most of the Renge schists have recorded green-
schist, epidote–blueschist or epidote–amphibolitefacies assemblages, and the lawsonite–blueschistfacies rocks are extremely rare in the Renge belt(Banno 1958; Nishimura et al. 1983; Nakamizu et al. 1989; Nishimura 1990). The Osayama
blueschists having the assemblages glaucophane +
lawsonite or glaucophane +pumpellyite belong to
a typical high-P/T type metamorphic facies seriesformed in the subduction zone. The differences ofthe recorded P/T conditions between the Osayamablueschists (typical high-P/T type) and the otherRenge schists (intermediate high-pressure type)may be due to the following reasons: (1) local diver-sity of geothermal gradient in the subduction zone;and/or (2) the different exhumation rate over-printed various lower P/T conditions. The Rengebasic schists of the Wakasa area have barroisiterimmed by actinolite, suggesting greenschist-facies overprint after epidote–amphibolite facies(T . Tsujimori, unpubl. data), although there is no evidence of greenschist-facies overprint in the200 T . Tsujimori and T . Itaya
Fig. 8 (a) Frequency distribution of K–Ar and Rb–Sr phengite (white mica) ages of Paleozoic high-pressure schists in southwestern Japa n (compiled
from Maruyama & Ueda 1974; Maruyama et al. 1978; Shibata & Ito 1978; Ueda et al. 1980; Shibata & Nishimura 1989; Isozaki et al. 1992; Kabashima
et al. 1995; Kunugiza et al. 1997), and of K–Ar hornblende ages for the amphibolites and gabbroic intrusions of the Oeyama ophiolite (Shibata et al.
1979; Shibata 1981; Nishimura & Shibata 1989; Nishina et al. 1990). The metamorphic facies of the schists are distinguished. PA, pumpellyite–actino-
lite facies; LBS, lawsonite–blueschist facies; EBS, epidote–blueschist facies; GS/EBS, transitional facies between greenschist and epidote–blueschist
facies; EA, epidote–amphibolite facies. (b) Isotopic age relations of the Cordilleran high-pressure metamorphic belts in southw estern Japan, western USA,
and eastern Australia (compiled from Patrick & Day 1995; Isozaki & Maruyama 1991; Little et al. 1993; Fukui et al. 1995; Nishimura 1998). The time
scale is after Harland et al. (1990).
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Paleozoic blueschist-facies metamorphism in SW Japan 201
Osayama blueschists. This suggests that the
Osayama blueschists were formed in the subduc-tion zone with the lowest geothermal gradient ofthe Renge belt. Absence of the greenschist faciesoverprinting in the Osayama blueschists suggestsa quick exhumation, which is supported by thehomogeneous K–Ar age, soon after the blueschist-facies metamorphism in the subduction zone.
Recent advanced studies of the Oeyama ophio-
lite have revealed that they represent a supra-subduction zone ophiolite formed beneath theback-arc basin or primitive arc setting (Arai &Yurimoto 1994, 1995). Some fragments of theOeyama ophiolite pre-dating the Osayama blue-schists have also suffered the blueschist faciesmetamorphism together with the Osayama blue-schists as mentioned before. It follows that theOeyama ophiolitic lithosphere was close to thetrench of the Carboniferous subduction zonesystem, and was eroded and dragged down to adeeper part of the subduction zone to undergoblueschist-facies metamorphism together with theRenge schists. Such tectonic erosion of the supra-subduction zone lithosphere has been documentedin the modern subduction system of the Marianaarc–trench system (Bloomer 1983; Maekawa et al.
1995).
COMPARISON WITH OTHER PALEOZOIC HIGH-P/T
SCHISTS IN SOUTHWESTERN JAPAN
The subduction-related metamorphic rocks pre-
dating the Renge schists in southwestern Japanhave already been reported (Fig. 8). In the Kuro-segawa belt of the Outer Zone of southwesternJapan, which is interpreted as a tectonic klippeconsisting of pre-Jurassic equivalents of the InnerZone (Isozaki & Itaya 1991), the pumpellyite–glaucophane schists, epidote–barroisite schistsand epidote–hornblende schist have been reported(Maruyama & Ueda 1974; Nakajima & Maruyama1978; Nakajima et al. 1978). The former has pet-
rographic features similar to the Osayamablueschists but gives K–Ar phengite ages of352–394 Ma (Ueda et al. 1980), significantly older
than those of the Osayama schists. The latter hastwo groups of K–Ar ages; one is 317–327 Ma (Uedaet al. 1980), similar to the Renge belt; and the
other is 402–445 Ma (Maruyama & Ueda 1974)demonstrating the oldest high-pressure schists insouthwestern Japan.
Amphibolites as tectonic blocks in the Oeyama
ophiolite have undergone epidote–amphibolitefacies metamorphism and have demonstratedhornblende K–Ar ages from 469 to 336 Ma
(Kurokawa 1985; Nishimura & Shibata 1989;Nishina et al. 1990). A garnet amphibolite giving a
K–Ar biotite age of 442 Ma (Matsumoto et al. 1981)
occurs as tectonic blocks within the 320-Ma Rengeschists in Hida Mountains (Nakamizu et al. 1989).
A clinopyroxene-bearing garnet–amphibolite witha K–Ar hornblende age of 409 Ma (Y oshikura et al.
1981) occurs as tectonic blocks in the Kurosegawabelt. To reveal the timing of igneous activity of theOeyama ophiolite, the gabbroic intrusions havebeen dated. They gave the hornblende K–Ar agesfrom 343 to 239 Ma (Shibata et al. 1979; Nishina
et al. 1990). Some of those ages are clearly younger
than those of the Renge schists, namely , they contradict the fact that some fragments of theOeyama ophiolite are enclosed as tectonic blocksin the serpentinite matrix and have suffered theRenge blueschist metamorphism together with the other blueschist blocks. These young ages ofthe gabbroic intrusions are likely to be due to therejuvenation by the post-dating metamorphismbecause the igneous brown hornblende in the gabbroic intrusions is commonly rimmed by actinolite (Y amaguchi 1989). Recently , Hayasaka et al. (1995) preliminarily reported the Sm–Nd
ages of ca560 Ma for gabbroic intrusions in central
Chugoku Mountains, suggesting a time of forma-tion of the Oeyama ophiolite as the Cambrian.
BLUESCHIST-FACIES METAMORPHISM DURING
PALEOZOIC OROGENY IN SOUTHWESTERN JAPAN
In the circum-Pacific orogenic belt, Paleozoic
blueschist facies metamorphic rocks also occur inwestern USA and eastern Australia (Fig. 8). TheSkookum Gulch blueschist ( ca 450 Ma) distributed
in the Yreka Terrane, eastern Klamath Mountains,is characterized by the mineral assemblage glaucophane +lawsonite, resembling the Osayama
blueschists (Cotkin 1987). The Skookum Gulchblueschist is tectonically overlain by serpentinizedperidotite of the Cambro-Ordovician Trinity ophio-lite, and blueschist contains 570-Ma tonalite blocksderived from Trinity ophiolite (Wallin et al. 1988).
The New England Fold Belt in eastern Australiaincludes three Paleozoic subduction-related metamorphic rocks, ca 260, ca 340–310, and ca
470 Ma (Fukui et al. 1995). The oldest rocks contain
467–481-Ma epidote–glaucophane schist occurringalong the ophiolitic serpentinite melange zone inthe Glenrock–Pigna Barney area, northeasternNew South Wales (Fukui et al. 1995). They are also
closely associated with 530-Ma ophiolitic rocks
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(Aitchison et al. 1992). It is considered that an active
continental margin formed an accretionarycomplex, high-P/T schists and volcano-plutonism atthe circum-Pacific orogenic belt (Isozaki 1996;Maruyama 1997). In southwestern Japan, theOrdovician schists of the Kurosegawa belt(Maruyama & Ueda 1974) is evidence for an incipi-ent subduction of the paleo-Pacific Plate, andtypical blueschists appear from the Devonian(Ueda et al. 1980) (Fig. 8). The petrologic and
geochronologic comparison of the Paleozoic high-P/T metamorphic rocks in southwestern Japanrevealed that the geothermal gradient in the sub-duction system was relatively high to form theepidote–amphibolite facies metamorphic rocks inLate Ordovician–Silurian time (e.g. Maruyama &Ueda 1974). The low geothermal gradient to formthe blueschists in the Devonian–Carboniferouscould be attained by an active subduction of thepaleo-Pacific oceanic plate. In early history of thecircum-Pacific orogenic belt, the subduction systemin western USA and eastern Australia had reachedthe low geothermal gradient to form the blueschistsas early as the Ordovician, when the system insouthwestern Japan still had a high gradient.
The tectonic association of Paleozoic ophiolite
and Paleozoic high-P/T schist is pervasivethroughout the circum-Pacific region. This sug-gests that each orogenic belt in the circum-Pacificregion had experienced an early history similar tothat in southwestern Japan.
ACKNOWLEDGEMENTS
We express sincere thanks to Drs T . Okada and
H. Takeshita for their help in age determination at Okayama University of Science. The seniorauthor is grateful to Dr A. Ishiwatari for daily discussion and the critical reading of the manu-script. Sincere thanks are due to Prof. S. Arai for his important suggestion about peridotitebodies in the Chugoku Mountains, and to Dr K.Kunugiza for discussion about metamorphic rocksin Hida Mountains. Prof. S. Banno and Prof. Y .Nishimura are thanked for their constructive criticisms. The present study was supported financially in part by JSPS Research Fellowshipsfor young scientists.
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Tsujimori (1999) - Blueschist_facies metamorphism during Paleozoic orogeny in southwestern Japan.txt
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VOLUME 91
THE JOURNAL OF GEOLOGY
May 1983
ORIGIN OF HYDROGEN AND CARBON DIOXIDE IN FAULT GASES AND ITS
RELATION TO FAULT ACTIVITY1
RYUICHI SUGISAKI, MASAHIKO IDO, HIROSHI TAKEDA, YUMIKO ISOBE, YOSHIMITSU HAYASHI,2
NORIAKI NAKAMURA,2 HIROSHI SATAKE,2 AND YOSHIHIKO MIZUTANI2
Department of Earth Sciences, Nagoya University, Nagoya 464, Japan
ABSTRACT
Gases occluded in fracture zones of active faults are characterized by a high concentration of H2 and/or
CO2. A predominant gas species-H2 or CO2-is related to the lithofacies which the fault cuts. Carbon
dioxide concentration in sediments fluctuates with temperature. This evidence and 13C of CO2 (about
-20%o) suggests that CO2 originates from organic materials. Carbon dioxide with 813C of -5%o to - 17%o
in brecciated gneiss containing marble may have been produced by interaction between organically derived
CO2 and the marble, or alternatively, may have been magmatically derived. Hydrogen usually occurs in
sheared silicate rocks, and its concentration fluctuates a great deal. The concentration of H2 from active
faults associated with historical earthquakes usually amounts to as high as several percent in maximum,
whereas the concentration of H2 from Quaternary faults not associated with historical earthquakes is at
most 100 ppm. Laboratory experiments showed that much H2 is generated from paste made of newly
pulverized rocks and water, suggesting that the fresh mineral surface formed by tectonic stresses reacts
with groundwater to produce H2. Since the mineral surface loses its activity with time, discrimination
between recently moved faults and other Quaternary faults can be made by the H2 concentration. Hydrogen
isotope thermometer, as well as field evidence, suggests a deep seated origin of H2 in an active fault.
Hydrogen measurements at monitoring stations offer information at depth on mechanisms that operate prior
to earthquakes.
INTRODUCTION
Active faults have occupied much attention
in society, as well as in the geosciences, be-
cause they are associated with earthquakes.
For example, when a power plant is con-
structed in an unstable area, activities of
faults in the area should be assessed and eval-
uated in order to take in earthquake design
for its stability.
Active faults have been studied mostly
from geological and geophysical viewpoints.
1 Manuscript received September 29, 1982; re-
vised December 30, 1982.
2 Department of Earth Sciences, Toyama Uni-
versity, Toyama 930, Japan.
[JOURNAL OF GEOLOGY, 1983, vol. 91, p. 239-258]
© 1983 by The University of Chicago. All rights
reserved.
0022-1376/83/9103-006$1.00 Recently, however, geochemical researches
on the problem have also been reported (e.g.,
King 1980); anomalies in composition of
gases such as Rn and He issuing from the
earth's interior have been considered as pos-
sible precursory signals for earthquakes.
These gases can possibly be expelled through
active faults, mineral springs, and other out-
lets. Sugisaki et al. (1980) classified the gases
along faults in central Japan into the follow-
ing two types: (1) Gases charged mainly
within fault gouges (Type I) are characterized
by a high concentration of H2 and CO2 and
lack of CH4. (2) Gases occurring as bubbles in
flowing mineral waters from fault zones
(Type II) show high concentration of He and
CH4; they generally lack CO2. Gases in
groundwater (Type II) have often been
studied for earthquake prediction after the
Tashkent earthquake of 1966 (Sadovsky et al.
239 NUMBER 3
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R. SUGISAKI ET AL.
FIG. 1.-The studied area (the right) and locations of the monitoring stations (A to F) and examined faults
(the left). The symbols of the fault in the figure are as follows: Ho: Horinji fault. Us: Ushikubi fault. Ato:
Atotsugawa fault. Kan: Kanbara fault. Kaz: Kazura fault. Hi: Hida fault. Ne: Neodani fault. Ate: Atera
fault. Mi: Misaka fault. Ch: Chobeizawa fault. Sa: Sanage Fault. Fu: Fukozu fault.
1972). Sugisaki (1978) pointed out the utility
of the He/Ar ratio in Type II gases as an
earthquake precursor. Furthermore, Sugisaki
(1981) found that a cyclic variation of the
He/Ar ratio in Type II gases was caused by
the tidal strain and suggested that deep-
seated gases are squeezed out by stresses
preceding an earthquake.
Type I gases can more easily be monitored
than Type II gases because the former can
be collected from wider sample localities.
Nevertheless, little information concerning
Type I gases is available. The present investi-
gation is focused on the CO2 and H2 of these
Type I gases charged in a fault zone. We ex-
amined, on the basis of field observation and
laboratory experiments, their spatial and
temporal variation, and discussed their origin
and the relationship between fault activities
and the compositional features of the gases.
The results of the investigation can be useful
for assessment and evaluation of fault activi-
ties and also perhaps for earthquake predic-
tion. PROCEDURES AND ANALYTICAL RESULTS
Sampling Location and Sampling Meth-
od.-Concentrations of H2 and CO2 fluctu-
ate a great deal, spatially and temporally,
in a fault. We tried to measure them as re-
peatedly and as widely as possible. Six moni-
toring stations were installed on five active
faults to examine the temporal variation in
gas composition. These faults are of large
scale and representative in central Japan.
Fault gases from some sites other than at the
monitoring stations in the five faults and
those from other faults were also examined.
The locations of the surveyed active faults
and the monitoring stations are shown in
figure 1. These faults show topographical and
geological evidence for fault movements dur-
ing the Quaternary.
Type I gases charged within a fracture zone
of a fault were collected by the method de-
scribed by Sugisaki et al. (1980): A gas sam-
pling plastic pipe was buried in a hole about 1
m deep within a fault zone. The gas in the
pipe was transferred into a sample flask. 240
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
Analytical Methods and Results.-Sample
gases were analyzed with a gas chromato-
graph apparatus, Ohkura Model 802 for H2,
He, Ne, Ar, CH4, N2 and 02. The mutual
separation of He, Ne, and H2 with gas ab-
sorption chromatography has so far hardly
been successful at room temperature. Re-
cently, we developed a simplified and accu-
rate determination of subsurface gases with a
gas chromatograph (Sugisaki et al. 1982). By
this method, we can accurately and rapidly
determine H2 in samples, the concentration
of which ranges from 0.1 ppm to several per-
cent. Oxygen was gas-chromatographically
determined by using argon as the carrier.
Carbon dioxide in some samples was deter-
mined with a gas detector tube in the field.
Error (2r) in the determination is almost 5%
for trace components (several ppm in concen-
tration) and 1% for other components.
Carbon dioxide for isotopic measurements
was purified in a trap cooled by liquid nitro-
gen, and the carbon isotope ratio was
measured with a mass spectrometer, Mi-
cromass Model 602E. The Sample for D/H
ratio measurements of H2 gas was prepared
by the same method as described by Kita et
al. (1980) and was measured by the mass
spectrometer. All results are reported as 8
values (in parts per thousand, %o) where
(%o) = (Rsample/Rstandard - 1) X 103
and R is the 13C/12C or D/H ratio. The stan-
dards are PDB (Cretaceous belemnite, Be-
lemnitella americana, from the Peedee for-
mation of South Carolina; Craig 1957) for
carbon and SMOW (standard Mean Ocean
Water; Craig 1961) for hydrogen. Analytical
errors are ±+0.1%o for 813C and ±+2%o for 8
D.
The analytical results were summarized in
tables 1, 2 and 3. Concentrations of CO2 and
H2 in these gases are higher than those in
atmospheric air, and they have large standard
deviations at each site, although N2 and Ar
concentrations do not fluctuate so much. This
shows that the fault gas is characterized by
both gases.
RELATION BETWEEN GAS
COMPOSITION AND GEOLOGY
Type I gas under consideration is usually
characterized by H2 and CO2 (Sugisaki et al. 1980), but these two species are not always
found together in a site on a fault zone. For
example, Wakita et al. (1980) observed H2
concentration as high as 3% at some sites on
the Yamazaki fault, which is a representative
active fault in southwest Japan. Yoshikawa et
al. (1982) reported a high concentration of
CO2, up to 12%, at a site in the Atera fault.
Sugisaki et al. (1980) observed CO2 of 1.05%
in a fault gouge of black schist along the Me-
dian Tectonic Line. The genetic relationship
between CO2 and H2 along fault zones, how-
ever, has been obscure. Our examination
conducted below was based on the expecta-
tion that a gas species emerging from a fault
zone depends on the geology of the area
which the fault cuts.
In order to examine the genetic relation-
ship between a gas and geology, we surveyed
along a transect across the Atera fault. This
fault can be geologically and geomorphically
traced more than 60 km from NW to SE in an
area of Cretaceous volcanics and granites. It
has been moving during the Quaternary
(Sugimura and Matsuda 1965). A monitoring
station at Tsukechi on the fault is shown as C
in figure 1. The area where the station is set
up is covered with terrace deposits from the
Pleistocene, which contain gravel, sand, and
a small amount of volcanic ash. The terrace is
underlain by the Nohi rhyolite welded tuff of
Cretaceous age. The welded tuff crops out on
the bed of a river which transects the fracture
zone of the fault (fig. 2). At this location, we
can examine the gas composition for two dif-
ferent strata appearing within the fracture
zone.
A series of plastic pipes for gas sampling
were put in shallow holes (about 1 m deep) on
the terrace; they were set by the Geological
Survey of Japan in July 1979 on a line across
the fracture zone. Maximum concentration of
CO2, up to 12%, was reported in the center of
the fracture zone when this monitoring site
was first installed (Yoshikawa et al. 1982).
Our measurements in December 1980 showed
a similar result (fig. 3), but no hydrogen was
detected in any pipe. The sample containing
much CO2 has little 02. At that time, we put
five pipes in the underlying welded tuff. The
vertical distance from the sampling pipes
settled in the terrace to the basement rocks is
at most 2 m. The spatial variation of CO2 and
H2 concentrations in the two strata is shown 241
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TABLE 1
RESULTS OF GAS ANALYSES AT THE MONITORING STATIONSa
Location A-i Measurement period
Jul.16,'81-Jan. 19,'82
Jul. 16,'81-Jan. 19,'82
Oct. 15,'81-Jul.21,'82
Oct. 15,'81-Jan. 19,'82
Nov. 12,'81-Jul.21,'82
Jul.16,'81-Nov.26,'81 Jul.16,'81-Nov.26,'81 Numbers of analyses 27 27 22 11 23 19 19 Gas
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
A-2 A-3 A-4 A-5 B-l B-2 Maximum 42 80.2 22.6 97
3.1 12 78.8 22.0 94 79
160 79.0 21.6 94
5.9 35 78.7 22.4 92 52
4900 94.5 22.3 1.10 3.7 2.6 79.6 20.7 94 1.00
200 83.0 21.8 97 11 Minimum 0
76.1 17.7 90 39 0
76.6 19.8 89 25 0
77.3 15.2 91 11 0
76.6 20.1 89 10 0
76.5 2.92 90 18 5
78.0 19.3 91 17 0
77.2 15.9 90 035 Average 4.0 77.6 20.4 92 1.1 1.3 77.8 20.9 92 44
8.7 78.1 19.9 92 1.2 3.9 77.6 21.3 91 26
757 84.1 13.4 99 94 1,3 78.7 20.1 92 45
21 78.6 20.4 93 062 Standard deviation 9.9 77 1.0 013 69
3.2 58 56 010 13
33 45 1.6 .0083 1.5 1.0 66 67 008 13
1270 4.4 4.6 047 47 1.2 83 73 015 25
47 1.2 1.2 014 019
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B-3 B-4 Jul. 16,'81-Nov.26,'81
Aug.6,'81-Jul.21,'82 Sep.15,'81-Jul.21,'82 Nov. 13,'81-Jan. 13,'82
May 27,'81-Apr.2,'82
May 27;'81-Mar.19,'82
May 27,'81-Mar.19,'82
May 27,'81-Mar.19,'82
B-5 D-1 E-1 20 26 16 11 9 9 9 9
E-2 H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
N2 (%)
02 (%)
Ar (%)
CO2 (%)
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%)
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%)
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%)
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%)
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%) 50 79.6 22.1 94 16
7600 80.6 21.0 96 59
11000 80.8 21.8 94 30
93570 6.2 19 79.7 21.1 97
10 5.8 24 92.3 19.3 1.10 10 5.8 20 82.3 19.8 1.01 12 6.2 21 89.7 17.5 1.08 63 6.0 20 81.7 18.5 1.01 0
76.9 19.4 90 074
2.2 77.7 18.3 90 03 0
76.7 18.2 90 045
354 4.8 18 76.7 11.2 89 9
5.1 17 84.3 8.2 95 9
5.1 19 79.3 16.7 51 7
5.4 17 81.6 9.2 93 1.0 4.8 17 80.0 17.2 92
E-3 2.7 78.3 20.7 92 12
1470 78.7 20.0 93 19
2310 78.8 19.9 92 11
32390 5.5 19 78.3 17.5 92
3.4 5.6 20 85.9 13.8 1.03 4.8 5.6 19 80.4 18.5 91
7.4 5.6 19 84.5 14.3 1.01 9.5 5.4 19 81.0 17.8 98 11 6.9 68 010 027
2040 70 65 012 13
3310 1.10 1.05 011 076
30910 4 42 99
3.1 025
3.8 25 1.9 3.7 2.9 043
2.7 23 59 97 91 15
4.0 25 1.2 2.6 2.9 041
20 31 1.0 54 45 028
(continued)
F-l
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R. SUGISAKI ET AL.
in figure 3. A sample from the basement con-
tained as much as 4150 ppm H2 but a small
amount of CO2. Samples containing much H2
showed concentrations of 02 lower than in
atmospheric air. He/Ar ratios of both strata
are substantially similar to atmospheric air.
From this field observation, we can draw
the following inference. The kind of gas
charged within a fracture zone is controlled
by the lithofacies which the fault cuts. Frac-
ture zones of sedimentary rocks are usually
associated with CO2, whereas, in igneous
rocks, H2 is usually found. Such a relation-
ship between gases and lithofacies is also ob-
served at monitoring station B, as stated
later.
The maximum concentration of these gases
occurs in the most intensively sheared part of
a fault zone, where fracturing of rocks is well
developed; the concentration on both sides of
the zone is lower (Sugisaki et al. 1980; Wakita
et al. 1980; Yoshikawa et al. 1982; and
others). This suggests that gases issue mainly
through the central zone of faults and that the
production of these gases is promoted by the
fault movement.
CARBON DIOXIDE
Spatial and Temporal Variation of Carbon
Dioxide along Faults.-The field observation
showed that CO2 tends to occur in sediments
and their related rocks. As far as our exami-
nation is concerned, the fault gas containing
more than 1% CO2 was found in the following
lithofacies: (1) black schist in the Sambagawa
metamorphic terrane (Sugisaki et al. 1980),
(2) terrace deposits and alluvial deposits, and
(3) biotite gneiss containing small amounts of
marble. In this paper, the cases of (2) and (3)
are described.
(i) Terrace and alluvial deposits: Moni-
toring station B is located on the Atotsugawa
fault, which lies about 70 km to the north of
the Atera fault, and trends NE-SW. This
fault runs more than 60 km along the Hida
plateau, a Tertiary raised peneplain. Along
the fault line, Paleozoic, Mesozoic, and
Cenozoic rocks and the Quaternary topo-
graphic surface are displaced right laterally
with the northern side upthrown (Matsuda
1966).
Monitoring station B was installed at Amo
on the Atotsugawa fault (fig. 1). At this sta-
tion, the fault cuts the gneiss and granite of 244
TABLE 1 (Continued)
Measurement period
May 27,'81-Mar.19,'82
May 27,'81-Mar. 19,'82
CO2 content at monitoring stations D, E and F is less than 0.05%. 02 content at monitoring stations A and B was calculated as the difference between 100% and the total amount of N2, Ar, and CO2.
Location F-2 F-3 Numbers of analyses 9 9 Gas
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%)
H2 (ppm)
He (ppm)
Ne (ppm)
N2 (%)
02 (%)
Ar (%) Maximum 50 6.9 22 96.9 14.8 1.14 26 8.8 24 98.5 15.3 1.16 Minimum 1
5.0 17 81.1 2.0 1.07 0.5 6.4 21 83.7 0.33 1.01 Average 5.6 6.0 20 90.8 7.0 1.09 4.4 8.0 22 96.0 4.02 1.11 Standard deviation 16 63 1.5 5.0 3.5 037
8.2 76 1.0 4.7 5.3 043
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
TABLE 2
CO2 AND ITS CARBON ISOTOPE RATIO IN SOME FAULT GASES
Location Date CO2(%) l13C(O/oo)
Monitoring station C-6 Aug. 9, '79 10.0 - 19.5
ML-4a Jul. 22, '79 1.05 -22.6
Monitoring station B-l Aug. 27, '81 .68 -23.0
B-3 Sep. 17, '81 .15 -21.5
B-4 Sep. 25, '81 .10 -21.7
aThis location was described in Figure 4 of Sugisaki et al. (1980).
TABLE 3
CO2 AND ITS CARBON ISOTOPE RATIO AT MONITORING STATION A
Sampling pipe Date CO2(O) b13C(O/oo)
A-1 Aug. 13, '81 1.7 - 8.5
A-2
A-3
A-4
A-5 Aug. 20, '81
Aug. 27, '81
Sep. 2, '81
Sep. 10, '81
Sep. 17, '81
Sep. 25, '81
Oct. 1, '81
Oct. 8, '81
Oct. 15, '81
Oct. 22, '81
Oct. 28, '81
Nov. 5, '81
Nov. 12, '81
Nov. 19, '81
Nov. 26, '81
Aug. 20, '81
Aug. 27, '81
Sep. 2, '81
Sep. 10, '81
Sep. 25, '81
Nov. 5, '81
May 11, '82
May 26, '82
Jun. 9, '82
Jun. 23, '82
Jul. 7, '82
Jul. 21, '82
Nov. 5, '82
May 11, '82
May 26, '82
Jun. 9, '82
Jun. 23, '82
Jul. 7, '82
Jul. 21, '82 2.1
1.5
2.0
1.3
.97
1.8
1.4
.83
1.1
.94
.60
.67
.50
.46
.53
.40
.43
.53
.55
.48
.73
1.04
.79
2.03
3.17
4.24
5.88
.37
.97
1.86
2.85
3.21
3.69
3.93 - 7.4
- 9.4
- 8.5
- 8.8
- 9.5
- 7.1
- 8.6
- 9.3
- 9.9
- 9.6
-10.8
-11.8
-11.6
-12.3
-11.1
-16.5
-16.6
-16.9
-17.4
-15.1
-11.1
-13.3
-13.8
-12.7
-12.6
-13.9
-13.8
-16.5
-12.0
- 8.3
- 7.6
- 6.5
- 6.3
- 5.7
the Hida metamorphic terrane. The fault
comprises five zones strongly sheared, each
of which is several meters wide. The total
width of the fracture zone amounts to 30 m,
and the strike of the fault is N75°E. Five sam-
pling pipes were put in a sheared zone, as
depicted in figure 4. One pipe (B-1) was put in the thin alluvial deposits of sand and clay.
Another (B-2) was put into the blue or black
gouge about 25 cm thick; the other three were
in the sheared gneiss and granite.
The gas composition has been continually
measured since July 1981, and the temporal
variation of CO2 concentration is shown in 245
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R. SUGISAKI ET AL.
N4
N5
FIG. 2.-Sketch map around monitoring station
C on the Atera fault. The lower figure is the section
on the line A-B of the upper one. Open circles with
a G number represent sampling pipes which were
put in the terrace deposits by Geological Survey of
Japan; closed circles with an N number are pipes
put in the basement rocks by tlhe present authors.
N1 N2 N3 NA N5
FIG. 3.-Lateral and vertical variation of H2,
CO2, and 02 concentrations in fault gases at moni-
toring station C. The upper and the lower figures
show the variation in terrace deposits (G: sampling
pipes in fig. 2) and in basement rocks (N: sampling
pipes), respectively. Open circles represent H2,
closed circles 02, and X's CO2. fault guage odori river
FIG. 4.-Sketch of cliff at monitoring station B
(Amo, Gifu) on the Atotsugawa fault, showing po-
sitions of sampling pipes (B-i to B-5).
figure 5. Carbon dioxide concentration in B-i
was higher than in any other pipe. The sec-
ond highest concentration was observed in B-
4, which was in the sheared granite but very
close to the alluvial deposits. Carbon dioxide
concentrations in other pipes do not exceed
0.2%. Another aspect observed is that CO2 in
B-i exceeded 1% in the summer, while in the
winter, it decreased to 0.2%. The meteoro-
logical data are shown in figure 5. This shows
a relationship between CO2 and the annual
temperature cycle.
(ii) Sheared gneiss containing marble:
Monitoring station A was installed at Ka-
megai on the Ushikubi fault located about
10 km to the north of the Atotsugawa fault.
The Ushikubi fault also runs more than 60 km
in the Hida plateau and is parallel to the Atot-
sugawa fault. The station was placed on the
fracture zone of granitic gneiss. On both sides
of the main fracture zone, there is a layer of
grey or black gouge about 5 cm thick, as out-
lined in figure 6. It is noted that thin lenses of
marble frequently occur on the right-hand
side of the fracture zone (biotite gneiss),
whereas, on the left-hand side (granitic
gneiss), there is no such occurrence. Five
sampling pipes were put in this station; one
pipe (A-5) in the fault gouge, another (A-3)
within the main fault zone.
As is the case with B-4, CO2 concentration
correlates with temperature at A-1 and A-3
(fig. 7). Another aspect observed at station A
is that the CO2 concentration increases from
the left granitic gneiss to the right biotite
gneiss (fig. 6. 6). Since marble lenses occur only
on the right-hand side of the outcrop, their
occurrence may be related to the higher CO2
concentration in the biotite gneiss. Attention
is directed to A-5, where much CO2 is found 246
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
B-1
8-2
B-3
B-4
B-5
JUL AUG SEP OCT NOV APR MAY JUN JUL
1981 1982
FIG. 5.-Temporal variation of CO2 at monitoring station B on the Atotsugawa fault, with the meteorolog-
ical data (atmospheric pressure, temperature, and precipitation).
together with much H2. A unique origin of the
CO2 in this case is implied, as stated later.
Carbon Isotopes.-The 613C values of CO2
in several samples are listed in tables 2 and 3.
They can be grouped into two clusters;
namely, about -20%o (table 2) and about
- 10%o (table 3). The former and the latter
are found in terrace and alluvial deposits (sta-
tions B and C), and in gneiss containing mar-
ble (station A), respectively. The temporal
variation of 8'3C at A-1 and A-5 in the latter group (station A) is correlated with that of
CO2 concentration (fig. 8), whereas the
isotopic ratios in the former group and in A-3
of the latter remain constant irrespective of
CO2 content.
Origin of C02.-The origin of CO2 charged
within fault zones deserves serious attention
for earthquake prediction, exemplified by Ir-
win and Barnes (1980) and Shapiro et al.
(1982).
The observations at Station C showed that 247
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R. SUGISAKI ET AL.
Fault gouge
Brecciated zone
Road
im
FIG. 6.-Sketch map around monitoring station
A (Kamegai, Toyama) on the Ushikubi fault, show-
ing positions of sampling pipes (A-i to A-5).
CO2 was hardly detected in the basement
rocks of rhyolitic tuff, and this suggests that
CO2 does not ascend through the fault from
depths but is produced in the terrace depos-
its. The CO2 in the terrace deposits is proba-
bly produced from organic materials by
biological processes. As shown in figure 5,
CO2 concentration in alluvial deposits (sta-
tion B) fluctuates with temperature. For a
given organism, its metabolic activity tends
to increase with temperature up to a certain
extent, and hence the production rate of CO2
in normal soil depends on its temperature.
Such a close relationship between CO2 pro-
duction and soil temperature has been re-
ported by many agronomists (e.g., Monteith
et al. 1964). The fact that CO2 production in
soil is followed by oxygen consumption has
also been known in the field of soil science.
Similarly, the lack of oxygen in gas samples
with higher CO2 was observed at station C
(fig. 3). Kanaori et al. (1982) reported that
there is also a reverse correlation between
CO2 and 02 in terrace deposits along the
Kaze fault.
Carbon isotope ratios constitute cogent
evidence for the origin of fault CO2. The 813C
value of about - 20%o for CO2 in terrace and
alluvial deposits, as listed in table 2, is com-
patible with CO2 production by the oxidation
of organic materials. Shapiro et al. (1982) also
interpreted that CO2 with - 22%o of o83C in
borehole air near the San Andreas fault origi-
nates from the oxidation of organic materials.
The origin of CO2 at monitoring station A is
perhaps explained as follows. The CO2 oc-
curs in association with much H2 in silicate
rocks, and therefore it cannot be regarded as simply biogenic. Their 13C value ranges from
-5 to -17%o. The fact that marble lenses
occur in the biotite gneiss only on the right
side of the main fracture zone suggests that
they are related to the genesis of CO2. As
stated earlier, CO2 is concentrated towards
the biotite gneiss. Similarly, the 13C value
increases towards the biotite gneiss. The s3C
value in such a range is observed in second-
ary calcites such as stalactites deposited from
fresh water (Gross 1964). They form when
carbonate in which calcites were dissolved is
precipitated from the groundwater. The car-
bon isotopic composition in groundwater is
determined by the isotopic composition of
the carbon from the two sources-carbonate
from limestone and CO2 from plants; their
8'3C values are approximately 0 and - 25%o,
respectively. Because marine carbonate is
dissolved by organically derived CO2, the
813C value of the secondary calcite clusters
around the average for the two sources,
namely -5 to -15%o (Sackett and Moore
1966). If the CO2 of fault gases is separated
from groundwater, a Y13C value similar to that
of the secondary calcite is predicted for this
CO2 gas. This process may be responsible for
the carbon isotope feature of the C02 in these
fault gases, because the 613C of a marble lens
from the outcrop of this fault is 0.5%o, which
is close to that of marine carbonates. If CO2 is
derived from bicarbonate which was pro-
duced by an interaction of organically de-
rived CO2 with marble, and it dissolves the
marble, then the 813C value of the last bicar-
bonate becomes higher. Thus, CO2 with a
813C value as high as -5%o could be gener-
ated.
An alternative explanation is that mag-
matic CO2 issues from depths through the
fault and simply is mixed with organically de-
rived C02; the 613C value of the magmatic
CO2 is generally accepted to be from 0 to
- 10%o (e.g., Hoefs 1973, p. 64). Since many
hydrothermal ore deposits and their presum-
ably related mineral springs occur along the
Ushikubi fault and the Atotsugawa fault, this
possibility cannot be excluded.
HYDROGEN
A significant amount of H2 occluded in
fault breccia and gouge was observed (Su-
gisaki et al. 1980; Wakita et al. 1980). As
shown by the observation at station C, H2 248
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
+ A-1
x A-2
o A-3
* A-4
* A-5
JUL AUG SEP OCT NOV DEC JAN APR MAY JUN JUL
1981 1982
FIG. 7.-Temporal variation of CO2 at monitoring station A on the Ushikubi fault, with the meteorolog-
ical data.
seemingly is produced in the fracture zone of
the basement volcanics. In the present study,
the behavior of H2 was examined in relation
to the other components. In order to better
understand the result, a series of controlled
laboratory experiments have been carried
out, and the H2 genesis discussed. Spatial Distribution of H2 and its Relation
to Fault Movements.-Data on H2 from ac-
tive faults have been accumulated by us as
well as by other investigators. Figure 9 sum-
marizes H2 concentrations in fault gases
(Type I). Within a fault, the H2 concentration
considerably varies because of complicated 249
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R. SUGISAKI ET AL.
1981 1982
FIG. 8.-Temporal variation of CO2 and '3C at
monitoring station A.
sampling conditions such as sampling time,
sampling site, sampling depth, and others.
Nevertheless, all the determined values over
1 ppm are plotted in the figure, indicating a
wide range in each fault. Faults are arranged in
order of the decreasing values of the high-
est concentration of H2 within each of
them. In spite of the complications in sam-
pling condition, the high concentration of H2
over several thousand ppm is found for all the
"historically active faults" except one (the
Hida fault). "Historically active fault" is
here designated as the fault associated with a
historic earthquake. The association is noted
in figure 9, although it cannot always be reli-
able for the earthquakes before earthquake
observations were performed scientifically.
H2 (ppm ) FAULT NAME EARTHQUAKE (M)
1 10 100 1000 10000 100000
FIG. 9.-H2 in active faults. The upper eight faults are associated with historical earthquakes except for
the Hida fault. Earthquakes associated with the lower seven faults are not known. The earthquake ages and
magnitudes were quoted from Okada and Ando (1980). Data for the Yamazaki fault, the Negoro fault, and
the Kaze fault were quoted from Wakita et al. (1980), Takehana et al. (1982), and Kanaori et al. (1982),
respectively. Other data were obtained by the present authors. Values below 1 ppm were not plotted. 250
On the other hand, the lower seven faults
in the figure are not associated with historic
earthquakes and are designated as "prehis-
torically active fault," although they have
sheared zones with gouges and breccias and
show topographical and geological evidence
for fault movements during the Quaternary.
For example, the fracture zones of the Mi-
saka fault (fig. 1) are clearly observed inside
the Enasan tunnel. The fault crosses the tun-
nel, and strongly sheared granites about 400
m wide crop out inside the tunnel. Three sam-
pling pipes, 2 m long each, were set with a
boring machine on the monitoring site (sta-
tion E), and the pipes were fixed with cement
during the tunneling work. Monitoring station
F inside the tunnel was also installed on the
Chobeizawa fault, and three pipes were set
up in the same way as at station E. We ex-
pected the best condition for sampling at sta-
tions E and F. The concentration of H2, how-
ever, was at most 70 ppm. The faults not
associated with historic earthquakes gener-
ally show a low concentration of H2. The dif-
ference in the highest H2 values between his-
torically active faults and prehistoric ones is
almost two orders of magnitude. This implies
that H2 is related to recent fault movements.
Temporal Variation of H2.-Hydrogen in
Type I gases fluctuates temporally and spa-
tially. For example, the concentration of H2
from pipes at monitoring stations A and B
ranges from 0 to 8000 ppm, although these
pipes are situated in the fracture zone only 10
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
JUL AUG SEP OCT NOV DEC JAN
1981 1982
FIG. 10a.-Temporal variation of H2 at moni-
toring station A.
m apart (fig. 10). Their temporal variation is
also significantly large. Hydrogen amounting
sometimes to several thousand ppm appears
in sites along historically active faults. Its ap-
pearances from each pipe at a station do not
temporally coincide with each other. On the
other hand, H2 concentration at monitoring
stations E and F along the prehistorically ac-
tive faults remains low, and does not fluc-
tuate so much (fig. 11).
What is the factor controlling the varia-
tion? We can hardly recognize a correlation
between H2 fluctuation and meteorological
data. From the evidence that a high concen-
tration of H2 is found for historically active
faults, and from laboratory experiments de-
scribed later, it is suggested that H2 produc-
tion is related to rock failure followed by
earthquakes. In the hope of carrying the sug-
gestion further, we plotted the total energy
released by earthquakes within 50 km from
monitoring station B (fig. 10). An association
of H2 variation with released energy, how-
ever, rarely emerges on the plot. This fact
will be discussed later.
Relationship of H2 to Other Gases.-
Successive measurements showed that a high JUL AUG SEP OCT NOV
1981
FIG. 10b.-Temporal variation of H2 at moni-
toring station B. The lower figure shows the daily
effective energy released by earthquakes, of which
epicentral distances are less than 50 km. The data
for earthquakes were observed at Kamitakara
Crustal Movement Observatory, Kyoto Universi-
ty: E (erg) was calculated from magnitude of earth-
quakes; r (km) is the epicentral distance from moni-
toring station B. The effective energy is repre-
sented by its averaged daily amount between two
collection times of sample gases.
concentration of H2 is followed by an oxygen
defect in fault gases. Figure 12 exemplifies
the relationship observed at monitoring sta-
tion D on the Atera fault, which is a histori-
cally active fault. The concentration of H2 is
inversely correlated with that of 02. Hydro-
gen is plotted against 02 in some faults (fig.
13). This indicates that points of historically
lJUN UL aUG SEP OCT NC DEC JAN FEB MAR
1981 1982
FIG. l.-Temporal variation of H2 at moni-
toring stations E (on the Misaka fault; closed cir-
cles) and F (on the Chobeizawa fault: open circles)
which are installed inside the Enasan tunnel. 251
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252 R. SUGISAKI ET Al
NOV DEC JAN
1981 1982
FIG. 12.-Temporal variation of H2 (open cir-
cles) and 02 (closed circles) at monitoring station D
on the Atera fault.
active faults are located on a narrow zone,
whereas those of prehistorically active faults
deviate from the zone to the lower left in
figure 13-gases short of 02 are characterized
by much H2 for historically active faults,
whereas, for the prehistorically active faults,
even a gas with a small amount of 02 contains
at most 100 ppm of H2. Using this plot of 02
versus H2 (fig. 13), as well as H2 concentra-
tion (fig. 9), we can discriminate between his-
torically active fault gases and prehistoric
ones.
Laboratory Experiments for H2 Produc-
tion.-Stress induced gas emission was ob-
served by Sugisaki (1981) in a field and was
experimentally studied by Giardini et al.
(1976) and Jiang and Li (1981). In order to
examine the genesis of H2 in fracture zones,
we carried out modeling experiments con-
trolled in the laboratory.
We focused on the experiments of rock-
water interaction. In a reaction flask, 250 ml
in volume, powder of rocks and minerals
ground finer than 60 mesh was reacted with
distilled water. Gases inside the flask were
drawn by a syringe and gas-chromatographi-
cally determined.
(i) Experiments with small amounts of ma-
terials: 10 g of rock powder and 5 ml distilled
water were placed in the reaction flask, and 02 (%)
FIG. 13.-H2 versus 02 plot in the fault gases.
Open circles represent samples from monitoring
stations E (on the Misaka fault) and F (on the
Chobeizawa fault) for prehistorically active faults.
Closed circles and X's represent samples from
monitoring station D (on the Atera fault) and the
Bodai pass (on the Negoro fault belonging to the
Median Tectonic Line), respectively; both faults
are designated as historically active faults. The
data of the Negoro fault are from Takehana et al.
(1982).
H2 inside the flask was successively deter-
mined. The experiment was conducted for
different rocks and minerals, under various
inside atmospheres and reaction tempera-
tures. Hydrogen was produced in all experi-
mental runs, but the production rate was dif-
ferent in each run (fig. 14). Blank tests using
only distilled water showed no H2 produc-
tion. The result indicated an increase in H2
production rate with increase in temperature.
The production rate of H2 when the flask was
filled with Ar gas was higher than that with 02
gas. Pegmatite powder produces more !H2
than does that of quartz and feldspar. In each
case, H2 increased linearly with time at an
earlier stage of the reaction, while at a later
stage, the production rate became lower.
(ii) Experiments at a lower gas/solid ratio:
Only 10 g of rock powder was used for the
above experiments, in which the H2 concen-
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
Pegmatite
0 25 50
Quartz
0 100 200
Feldspar
0 100 200
Hour
FIG. 14.-Experimental results of H2 production
for three silicate powders (pegmatite, quartz, and
feldspar), different reaction temperatures (35°C
and 700C), and two atmospheres inside the reaction
flask (Ar and 02). 10 g silicate powder is reacted
with 5 ml distilled water in the reaction flask.
tration in the flask was at most 600 ppm. In
reality, however, the fracture zone contains
only small amount of gases. The ratio of gas
to solid is very low in fault breccia and gouge.
Is it possible to obtain H2 concentration as
high as several percent found in an histori-
cally active fault, if we make the gas/solid
ratio much lower in the experiment? We con-
ducted another experiment under a circum-
stance similar to fracture zones: We mea-
sured 02 this time, as well as H2. In the flask,
we placed 300 g of rock powder and mixed it
with 80 g of water to make a thick paste about
200 ml in volume. The volume of gas phase
left in the upper part of the flask containing 253
the rock paste was about 50 ml. This remain-
ing space was not replaced by a pure gas but
was filled with the room atmospheric air. The
flask was kept at 50°C. The various powders
were used for the experiments as follows: (1)
Pegmatite. (2) Granite. (3) Biotite extracted
from the granite; the purity was about 70%.
(4) The rest of the granite after the biotite was
extracted, which consisted mostly of quartz
and feldspar. (5) Fault gouge of granite at
monitoring station E inside the Enasan tun-
nel. (6) Fault gouge of granite at monitoring
station D on the Atera fault. (7) Precipitated
silica.
The experimental results in figure 15 indi-
cate a decrease in 02 for all cases and a
significant difference in the amount of pro-
duced H2 in each run. The concentration of
H2 for pegmatite amounted to a maximum of
3%. The second highest production of H2 was
shown by biotite. The residue of the granite
after the biotite was extracted has a H2 pro-
ducing ability lower than the whole granite.
This indicates that the H2 production rate for
biotite is higher than that for feldspar and
quartz. This is also suggested by the experi-
ments with small amounts of materials, as de-
scribed before. On the other hand, the silica
precipitated from water glass solution pro-
duced at most 10 ppm H2. The gouges from
the two active faults also hardly generated
H2.
The relationship between 02 and H2 in
these experiments is shown in figure 16.
There exist two trends in the figure. One is
the trend of much H2 production, the other,
the trend of 02 deficiency with a little H2 pro-
duction. On the plot of 02 versus H2, the for-
mer trend approximately corresponds to that
for historically active faults, whereas the lat-
ter corresponds to prehistoric ones (fig. 13).
The reaction flask is not completely air-
tight. Therefore, some amount of H2 might
escape out of the flask and 02 might get into
the vessel. Thus, the value determined for H2
production and 02 consumption in these ex-
periments is regarded as minimum.
Discussion on the H2 Generation in
Faults.-Molecular H2 can be formed by
various processes in nature. For example,
biological generation of H2 gases has been re-
ported by several authors (e.g., Koyama
1963). Hydrogen produced by a biological
process, however, is associated with CO2 and
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* Pegmatite
+ Biotite separated from granite
O Granite
e Felsic residue after biotite was
separated from granite
D Gouge from the Atera fault
X Precipitated silica (reagent)
Q Gouge from the Misaka fault
Hour
FIG. 15.-Experimental results for H2 production and 02 consumption. 300 g silicate powder and 80 ml
distilled water are mixed and emplaced in the reaction flask. Upper gas phase of the flask is atmospheric air.
Reaction temperature is 50°C.
CH4. Fault gases with much H2 occurring in
igneous rocks do not contain CH4; CO2 is
present in trace amounts, if any. Accord-
ingly, H2 of fault gases cannot be considered
as biogenic. Therefore, we suggest that H2 is
produced by an inorganic process. A series of
laboratory experiments carried out in the
present study support the inorganic genesis of H2. The materials used for the experiments
were silicate rocks and minerals, and biolog-
ical processes cannot occur in the H2 produc-
tion. Wakita et al. (1980) invoked the study of
Schrader et al. (1969) to explain the genera-
tion of H2 along the Yamazaki fault. Accord-
ing to Schrader et al., quartz will be trans-
formed into the amorphous state during R. SUGISAKI ET AL. 254
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
02(%)
FIG. 16.-H2 versus 02 plot in the gases of ex-
periments shown in figure 15. The symbols are the
same as those in figure 15.
vibration-milling, and radicals arise on the
surface of quartz. These radicals react with
water to produce H2 as follows:
-Si- + H20 -* Si - OH + H-
2H- -- H2
The materials used for the experiments
were newly pulverized. Hydrogen was pre-
sumably produced by the same mechanism as
that claimed by Schrader et al. Schrader et al.
studied only quartz, but we showed that
feldspar and other silicate minerals also pro-
duce H2 in the present experiments. Among
these minerals, biotite has sheet structure. It
seems likely that biotite powder has a large
surface area with radicals and that ferrous
iron in biotite may promote the reaction,
hence biotite shows higher H2 production.
The production rate observed in the experi-
ments decreases with time. This may be as-
cribed to the consumption of radicals on the mineral surface. The H2 produced by radicals
with high activity may directly react with 02,
resulting in the decrease of 02 in the gas
phase.
The most satisfactory explanation as to the
origin of H2 along active faults is offered on
the basis of field observations, laboratory ex-
periments, and the hypothesis of the active
mineral surface described above. Through
fault movements, crustal rocks deform and
eventually break; at the same time, cracks
open and extend. Fresh mineral surface
formed in the progress of the tectonic stress
may react with groundwater which flows into
the cracks. This interaction produces H2;
meanwhile, 02, if it is present there, is con-
sumed. As the failure of silicate rocks ad-
vances, fresh surfaces greatly increase and
produce much H2. Therefore, a lot of H2 is
associated with a recently moved fault. The
concentration of H2 thus produced in the
fracture zone (Type I gases) sometimes
amounts to several percent.
The fault movements regarded as responsible
for the H2 production in this paper, however,
may not always be related with the main
shock of a historic earthquake. Along the his-
torically active faults listed in figure 9, many
aftershocks and microearthquakes are occur-
ring even now. It is inferred that these minor
activities are related to H2 production, be-
cause the decrease of H2 concentration does
not seem to coincide with the elapse of time
after the main shocks, and moreover, a high
concentration of H2 appears irregularly in an
active fault, as observed at monitoring sta-
tions A and B.
On the other hand, Giardini et al. (1976)
suggested an inorganic origin of H2 through
experiments in which H2, CH4, CO, N2, and
CO2 are released when igneous rock speci-
mess are crushed at room temperature in a
vacuum. Our field observations, however,
show that H2 in Type I gases is not associated
with CO2 and CH4. Hence, the gases oc-
cluded in rocks cannot be considered as a
main source of H2 issuing from faults, al-
though the gases may sometimes be released
by rock failure, thus supplying small amounts
of H2.
The reaction with water consumes radicals
on mineral surfaces, and accordingly the H2
producing ability declines, as seen in our ex-
periments, in which H2 production rate de-
creases with time. As fault activities de- 255
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R. SUGISAKI ET AL.
crease, the H2 production drops. As a result,
the historically active fault can be dis-
criminated from the prehistoric one with re-
spect to the H2 amount and the relationship of
H2 versus 02 (figs. 9 and 12). For example,
the fault gouge from monitoring station E (in-
side the Enasan tunnel) has hardly any pro-
ducing ability of H2 (fig. 15); this station is
located on the Misaka fault, which is neither
associated with historical earthquakes nor
microearthquakes, and shows a H2 concen-
tration of 70 ppm at most (fig. 11).
Kanaori et al. (1980) observed that the sur-
face texture of quartz grains of fault gouges
showed various patterns, and they claimed
that quartz grains are progressively corroded
by groundwater after faulting. This mor-
phological change of mineral grains is cer-
tainly related to rock-water interaction in the
fault gouge.
The fault gouge at monitoring station D
located on the historical Atera fault also does
not have the H2 producing ability of the labo-
ratory experiment (fig. 15), although several
percent H2 are observed at this station (fig.
12). This gouge was taken from an outcrop
near the Earth's surface: it may have been in
contact with meteoric water for a long time,
and have lost the H2 producing ability. This
suggests the deep-seated origin of H2 of high
concentration along the fault.
Hydrogen isotopic evidence reinforces the
above inference. If we assume that hydrogen
isotope equilibrium between water and hy-
drogen gas at depth remains unchanged while
the gas rises up to the surface, we can calcu-
late the equilibrium temperature, from the D/
H ratio of the hydrogen gas and the ground-
water, using the fractionation factor for
hydrogen isotope exchange between liquid
water and hydrogen gas. We adopted the pro-
cedure reported by Arnason (1977), although,
for the calculation, we used new data con-
cerning fractionation factors published by
Richet et al. (1977). The 8 D values of H2
at station D and groundwater near the out-
crop were determined to be -675%o and
-55.9%o, respectively. With these data and
the fractionation factor, the temperature of
isotopic exchange equilibrium was calculated
to be 73°C. If we assume the geothermal gra-
dient in the area to be 2°C/100 m-3°C/100
m, we obtain 2.9 km-1.9 km for the depth
at which the equilibrium reaction was
quenched. Kita et al. (1980) calculated the depth of the reaction along the Yamazaki
fault with a similar method, and they found
that there was an agreement between the
isotopically estimated depth and the foci of
microearthquakes along the fault.
Whether or not all the H2 observed in frac-
ture zones stems from such a deep location is
not yet certain, but it certainly takes some
time for the H2 gas produced at depth to
travel through fissures and reach the surface.
Hydrogen will not be produced soon after the
rock failure but will be generated after the
groundwater permeates into newly formed
cracks. Furthermore, the H2 molecule is
chemically active, and it might be involved in
a reaction with wall rocks on the way of its
ascent. These complicated processes can be
responsible for the irregular emanation of H2
which was observed at monitoring station B
(fig. 10b), since it is unlikely to be related to
the energy released by earthquakes.
CONCLUDING REMARKS
"What are useful geochemical precur-
sors?" is a significant problem in seismo-
geochemistry. At present the geochemical
parameter mostly studied in the world is Rn.
Sugisaki (1978, 1981) pointed out the utility of
the He/Ar ratio. Carbon dioxide and hydro-
gen of fault gases described in the present
investigation are worth studying.
Carbon dioxide related to faults abundantly
occurs in alluvial and terrace deposits, but it
is mostly biogenic, in view of isotopic data
and other evidence. It is likely that, as a re-
sult of faulting, circulation of waters and soil
airs along fracture zones is promoted, and
biological activity in them increases. Carbon
dioxide of this kind does not stem from
depths and, accordingly, we cannot expect it
to be a useful precursor to earthquakes, al-
though CO2 can be applied to the prospect of
a fracture zone buried under sediments. In
contrast, CO2 at monitoring station A might
be of magmatic origin, as stated earlier. Such
deep-seated CO2 might be a useful precursor
for the earthquake prediction, as suggested
by Irwin and Barnes (1980).
The present experiments showed that
much H2 is produced by the reaction of pul-
verized rock with water. The production of
H2 in Type I gases can be attributed to this
reaction. The fracturing of rocks at depth
must usually occur from crustal stresses re-
sulting from earthquakes. Prior to earth- 256
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HYDROGEN AND CARBON DIOXIDE IN FAULT GASES
quakes, stresses increase and cracks of crus-
tal rocks are newly formed; groundwaters
flow into the newly created pores (e.g.,
Scholz et al. 1973). Hydrogen will be pro-
duced after this stage. The evidence of the
experiments and hydrogen isotopic composi-
tion is compatible with the deep-seated ori-
gin of H2. Therefore, H2 in fault gases will
provide information concerning stresses at
depth.
As stated earlier, the H2 fluctuation at sta-
tion B seemingly is not related to the energy
released by earthquakes. The largest mag-
nitude of earthquakes that occurred during
the observation period at the station was 3.1
(August 10, 1981): its epicentral distance was
37 km. For the distance, an earthquake with
magnitude larger than 3.5 will give Rn sig-
nals (Fleischer 1981) and He/Ar ratio anom-
aly (Sugisaki 1978, Sugisaki 1980). By the
same token, if larger earthquakes occur, H2
changes associated with them are expected to
be observed.
Hydrogen is characteristic of Type I gases
(fault airs), but it sometimes appears in Type
II gases (bubbles from mineral springs). For
example, at two monitoring stations of Type
II gases, concentration of H2 from 10 to 400
ppm appeared in bubble gases during about
one month prior to three earthquakes in cen-
tral Japan on August 11, 15, and 18, 1981.
Their magnitudes were 4.0, 4.9, and 5.1, re-
spectively. The epicentral distances were
about 50 km for each. Hydrogen is not usu-
ally observed at these stations. The period of
Hz emergence approximately coincided with
that of He/Ar ratio anomaly (Sugisaki et al.
1981). This example of a connection between
seismic activities and H2 emergence in
groundwater suggests that H2 measurements
at monitoring stations might give information
on mechanisms operative prior to earth-
quakes.
The present investigation was carried out
within a limited area and time. Much more
data must be accumulated before the useful-
ness of these gases for earthquake prediction
is confirmed.
Besides earthquake prediction, chemical
features of fault gases, especially H2, would
be useful for identifying recently active
faults. It is difficult to estimate whether a
fault is active or inactive, or merely dormant.
The term "active fault" is ambiguous, but in
a rough way, an active fault is defined as one that has moved in the recent past, and the
geological and seismological evidence sug-
gests that it may move in the near future. In
order to recognize active faults, it is neces-
sary to determine if movement has taken
place on faults and, if so, to date this move-
ment. Several methods hitherto have been
used for the examination, for example, field
mapping, trenching and the use of radiocar-
bon dating, geodetic measurements, aerial
photography, and others.
Present investigation about H2 from faults
may add another new method to those men-
tioned above. It shows that H2 as high as sev-
eral percent in amount occurs along some
faults of central Japan. It is evident that these
faults have moved in the historical period,
whereas some other faults which are known
not to have moved in the recent past do not
discharge much H2. If this correlation be-
tween H2 concentration and fault movement
can be applied to faults of other regions, it
may prove to be a powerful means for iden-
tifying active faults.
ACKNOWLEDGMENT.-The authors wish to
thank Dr. T. Kuwahara, Meijo University,
Drs. T. Ui and S. Fujii, Toyama University
and Dr. M. Adachi, Nagoya University for
their guidance in field operation and provid-
ing geological information. Thanks are also
due to Drs. H. Wada and M. Doi, Kyoto Uni-
versity, and the staffs of Toyama and Taka-
yama Local Meteorological Observatories for
providing the seismological and meteorolog-
ical data. Special thanks go to Dr. N. Tono
and his colleagues, Geological Survey of Ja-
pan who kindly gave permission to use the
sampling pipes at monitoring station C. Dr. S.
Matsuo, Tokyo Institute of Technology, pro-
vided the facilities for the purification of hy-
drogen gas in the fault gases for deuterium
analyses. The authors are grateful to Dr. M.
Kusakabe, I. Kawasaki and A. Takeuchi,
Toyama University for many stimulating dis-
cussions about the problem involved. The co-
operations of Mr. S. Yoshioka, Nagoya Uni-
versity for his technical assistance and Miss
H. Izumi, Toyama University for her assis-
tance in collecting sample were invaluable.
The authors are indebted to Dr. S. Oana for
reading the manuscript. This work was sup-
ported in part by Grant from the Ministry of
Education, Science and Culture, Japan. 257
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R. SUGISAKI ET AL.
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NAKAMURA, Y., 1982, Emission of hydrogen
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© 2011 Geological Society of AmericaV olcano-tectonic interactions during rapid plate-boundary evolution in
the Kyushu region, SW Japan
S.H. Mahony1,†, L.M. Wallace2, M. Miyoshi3, P. Villamor2, R.S.J. Sparks1, and T. Hasenaka4
1Department of Earth Sciences, University of Bristol, Wills Memorial Building, Queens Road, Bristol, BS8 1RJ, UK
2GNS Science, 1 Fairway Drive, Avalon, Lower Hutt 5010, New Zealand
3Beppu Geothermal Research Laboratory, Institute for Geothermal Sciences, Kyoto University, Noguchibaru,
Beppu, Oita 874-0903, Japan
4Department of Earth and Environmental Sciences, Kumamoto University, Kumamoto 860-8555, Japan
2201GSA Bulletin ; November/December 2011; v. 123; no. 11/12; p. 2201–2223; doi: 10.1130/B30408.1; 14 fi gures; 4 tables; Data Repository item 2011289.†E-mail: sue.mahony@bristol.ac.ukABSTRACT
Evolution of the local plate tectonic and
volcanic system relationship at Kyushu Island is defi ned by major changes in tec-
tonics and volcanic style at ca. 15, 10, 6, and 2 Ma. Plate reconstructions presented here suggest that prior to 15 Ma, the Pacifi c
plate subduction dominated Kyushu tecton-ics. From 15 to 6 Ma, the evolving relative plate motions shifted the triple junction between the Pacifi c plate, Philippine Sea
plate, and southwest Japan northwards, so that the Philippine Sea plate was sub-ducted beneath Kyushu. We suggest that a lack of subduction-related volcanism from 10 to 6 Ma is due to shallow subduction of the young Shikoku Basin lithosphere. By 6–5 Ma, changes in the Philippine Sea plate motion led to more rapid, nearly trench-normal, subduction of the Eocene west Phil-ippine Basin crust beneath Kyushu. This model is supported by an increase in arc-like geochemistry of lavas since ca. 6.5 Ma. Subduction of fl uid-rich features such as the
Kyushu-Palau ridge introduced large vol-umes of fl uids into the Kyushu arc system,
leading to voluminous volcanism across Kyushu, focused particularly in areas where the ridge subduction occurs in tandem with local extensional tectonics. Key issues, such as the timing of Izu arc collision with central Japan and the history of motion of the Phil-ippine Sea plate, are reassessed here, result-ing in a model that favors Izu arc–central Japan collision at ca. 8–6 Ma, rather than the more widely accepted date of ca. 15 Ma.INTRODUCTION
At subduction margins, volcanic and tectonic
processes are intrinsically linked. The distri-bution and style of volcanism depends on the regional tectonic framework, factors includ-ing: subducting slab dynamics, morphology and composition, mantle-wedge geodynam-ics, upper-plate structure and composition, and regional tectonic deformation. Likewise, tecton-ics can be strongly infl uenced by the location of
active volcanism; for example, the upper plate can be thermally weakened by magmatism, and thereby magmatism can lead to the localization of deformation. V olcanic processes can perturb regional stress fi elds and infl uence the style of
faulting. Some plate boundaries undergo rapid temporal and spatial changes, particularly where there is strong along-strike variation in the age of the subducting plate, or where a triple junction is present.
Kyushu is located in southwest Japan
(Fig. 1), where the Philippine Sea plate is sub-ducted beneath the Amurian plate, a subplate of the Eurasian plate (e.g., Petit and Fournier, 2005; DeMets et al., 2010). Philippine Sea plate subduction has led to a well-developed volcanic arc in Kyushu. However, over the past 15 Ma, there have been periods of apparent slow sub-duction, changes in subduction direction, sub-duction of a ridge, shallow plate subduction, clockwise and anticlockwise vertical axis rota-tion of the forearc region between the volcanic front and the trench, rollback of the subducting slab and backarc and intra-arc rifting, as well as changes in relative motion of the adjacent tec-tonic plates. These major Cenozoic changes in tectonic setting (e.g., Cambray and Cadet, 1994; Kamata and Kodama, 1994; Yamaji, 2003) have led to a complex volcanic history; since 15 Ma
there have been “fl ood” basalt and andesite
lavas, backarc monogenetic volcanoes, and arc volcanoes varying in behavior from lava domes to calderas.
Abundant geophysical and geological infor-
mation for Kyushu allows an in-depth study of how volcanism over the past 15 Ma has responded to changes in plate convergence directions and subducting slab characteristics. In Kyushu the volcanic-tectonic system is young, complex, and rapidly changing. Geochemical data are integrated into our analysis by using patterns in distribution of fl uid mobile elements,
giving further insight into the volcanic and tec-tonic evolution. We suggest that the volcano-tectonic evolution of the Kyushu region is con-sistent with the history of Philippine Sea plate motion as constrained by paleomagnetic and seafl oor spreading studies (Hall et al., 1995a,
1995b; Sdrolias et al., 2004). Our hypothesis for the volcano-tectonic evolution of Kyushu is at odds with most published interpretations by suggesting a mid-Miocene onset of the Izu arc collision with central Japan (e.g., Seno and Maruyama, 1984; Watanabe, 2005; Saito et al., 2007). This highlights the need to resolve the fundamental disconnect between current under-standing of past Philippine Sea plate motion and the timing of Izu arc collision.
CENOZOIC TECTONIC SETTING IN
THE REGION OF SOUTHWEST JAPAN
Throughout the Cenozoic, the long-term evo-
lution of the Pacifi c margin of Southeast Asia
provided the overriding driving force behind more localized margin changes in Japan. Prior to late Cenozoic time, the entire length of the
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Mahony et al.
2202 Geological Society of America Bulletin, November/December 2011Japanese islands was underthrust by the Pacifi c
plate (Taira, 2001; Hall, 2002). Since that time, paleomagnetic data and spreading histories of marginal basins have been interpreted to sug-gest three major periods of regional tectonic development in the adjacent Southeast Asia and the southwest Pacifi c regions, at 45 Ma, 25 Ma,
and 5 Ma (Hall et al., 1995a, 1995b; Hall, 1998).
Although the exact history and timing is con-
troversial, the Sea of Japan opened up west of Japan sometime between the mid-Oligocene and the mid-Miocene; some data sets indicate a 12–30 Ma timing (Otofuji and Matsuda, 1983;
Tamaki et al., 1992), while other data sets sug-gest a more short-lived opening history from 14 to 16 Ma (Otofuji et al., 1991). To accom-modate the Sea of Japan opening, the Japanese islands parted from the Asian mainland (Tamaki et al., 1992; Jolivet et al., 1994; Maruyama et al., 1997; Lee et al., 1999), with ~45° clockwise rotation of Kyushu and southwest Honshu and ~45° anticlockwise rotation of northeast Hon-shu and Hokkaido (Jolivet and Tamaki, 1992; Otofuji et al., 1991, 1994). The Shikoku Basin
Sea of Japan
Ryukyu
TrenchNankai Trough
KPR
Okinawa Trough
Parece Vela
BasinJapan TrenchKuril Trench
Izu-Bonin-M
ar
ianas TrenchIzu-Bonin-Marianas Arc
Shikoku Basin
(26–15 Ma)
West Philippine
Basin (60–40 Ma)90 mm/yr
72 mm/yr
79 mm/yrOkhotskplate
24° 28° 32° 36° 40° 44°
Pacificplate
130 MaKYUSHUEurasian
plate
(Amurian subplate)
10 mm/yr124° 128° 132° 136° 140° 144° 148°
Philippine Sea plate HONSHUHOKKAIDO
SHIKOKU
N0 150 km
Figure 1. Overview of plate tectonics of the Japanese arc system, indicating the interaction
between the three main plates—the Eurasian plate (Amurian subplate), the Philippine Sea plate, and the Pacifi c plate. White outlined areas indicate land. Convergence rates at each
respective plate boundary are labeled in mm/yr. The gray lines with solid triangles mark subduction zones. The gray dashed line cutting across Honshu and up the western side of Honshu and Hokkaido marks a zone of convergent tectonics, which is not a subduction zone. The main subduction zones are marked—the Ryukyu Trench, Nankai Trough, Izu-Bonin-Mariana Trench, Japan Trench, and Kuril Trench. The Okinawa Trough is a recent backarc basin (since 6 Ma). The Sea of Japan, West Philippine Basin, Shikoku Basin, and Parece Vela Basin are mature basins. The Kyushu-Palau ridge (KPR) is a remnant arc currently being subducted beneath Kyushu. Bathymetry image © 2009 Terrametrics.lithosphere formed due to rifting and spreading
in the eastern part of the Philippine Sea plate between 30 and 15 Ma (Watts and Weissel, 1975; Kobayashi and Nakada, 1979; Okino et al., 1999; Hall, 2002; Sdrolias et al., 2004) (Fig. 1). The Kyushu-Palau ridge divided the young Shi-koku Basin (eastern Philippine Sea plate) from the 40–60 Ma West Philippine Basin (western Philippine Sea plate) lithosphere. The Kyushu-Palau ridge is a remnant Eocene– Oligocene arc that split away from the Izu-Bonin-Mariana arc during Shikoku Basin rifting. These events strongly infl uenced the position and nature of
the triple junction (Fig. 1) between the Pacifi c
plate, Philippine Sea plate, and Eurasian plate (or Amurian subplate), which has evolved since 15 Ma, impacting the tectonic history of central and southwest Japan.
Present Day
Japan is currently located astride two subduc-
tion margins: the Pacifi c plate is being subducted
beneath northern Japan, while the Philippine Sea plate is being subducted beneath southwest Japan (Fig. 1). Much of the upper plate in south-west Japan is part of the Amurian plate (Wei and Seno, 1998). The contemporary Philippine Sea plate adjacent to southwest Japan can be divided spatially into two parts due to lithospheric age differences. Young buoyant Shikoku Basin lith-osphere (ca. 26–15 Ma) is subducted beneath southwest Honshu at the Nankai Trough, while the older (40–60 Ma) West Philippine Basin lithosphere is subducted beneath central and southern Kyushu at the Ryukyu Trench. A vol-canic arc has not yet developed in southwest Honshu where the Shikoku Basin is subducted at a shallow angle and the leading edge has only reached 70 km depth (Fig. 2); in contrast the West Philippine Basin portion of the Philippine Sea plate is subducted seismically to depths of 150–200 km beneath Kyushu (Fig. 2), and a well-developed volcanic arc is located above the 105 km depth contour (England et al., 2004; Nakajima and Hasegawa, 2007).
Major Active Tectonic Features in Kyushu
Major tectonic features in Kyushu are two
extensional grabens and two strike-slip domains. In central Kyushu, the Beppu-Shimabara graben (Fig. 3) exhibits ongoing north-south exten-sional faulting (e.g., Kamata, 1992; Kamata and Kodama, 1999), which coincides with abundant magmatic activity. The extensional faulting con-tinues southwest to backarc rifting in the Oki-nawa Trough (e.g., Lee et al., 1980; Sibuet et al., 1987). The Kagoshima graben in southern Kyushu (Fig. 3) hosts recent active east-west
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2203
130° 131° 132° 133° 134° 135° 136° 137° 138° 139° 140°E
37°
36° 35° 34° 33° 32°31°N
0 100 kmN
extension and active normal faulting (e.g., Ara-
maki, 1984; Kodama et al., 1995; AIST, 2008) at rates up to 7–8 mm/yr (Wallace et al., 2009a), which may be related to the recent trenchward migration of the active volcanic arc due to slab rollback (Yamaji, 2003). The southern bound-ary of Beppu-Shimabara graben extension in Kyushu is the Oita-Kumamoto line, a system of dextral strike-slip faults that are thought to be the westward continuation of Shikoku Island’s Median Tectonic Line (Okamura et al., 1995; Fig. 3). An east-west–trending dis-continuity in global positioning system (GPS) velocities across southern Kyushu, just north of the Kagoshima graben, indicates active left-lateral shear cutting across southern Kyushu (Nishimura and Hashimoto, 2006; Wallace et al., 2009a). Wallace et al. (2009a) have sug-gested that this shear occurs in response to sub-duction of the Kyushu-Palau ridge.
VOLCANIC HISTORY OF KYUSHU
This paper discusses volcanism from 15 to
10 Ma, 10–6 Ma, 6–2 Ma, and Quaternary vol-canism from 2 to 0 Ma. These time divisions are based on major changes in the style and intensity Figure 2. Confi guration of the subducted Philippine Sea plate modifi ed from Nakajima and Hasegawa
(2007). Triangles mark the locations of active volcanoes. The thick gray dashed line marks the position of the leading edge of the subducted Philippine Sea plate. The gray contour lines mark the depth contours of the subducted Philippine Sea plate slab. The tight contours over Kyushu represent the steeply dipping West Philippine basin lithosphere, compared to the shallower dip of the Shikoku Basin lithosphere under southwest Honshu.
1GSA Data Repository item 2011289, Geochemi-
cal discrimination diagrams of the volcanic rocks in Kyushu, is available at http://www.geosociety.org/pubs/ft2011.htm or by request to editing@ geosociety.org.of volcanism, geochemical characteristics of mag-
mas, and the presence of volcanic gaps. Detailed volcanic histories are presented in Tables 1–4, which describe the general changes within each time division. Unless otherwise stated in the appropriate table, all information regarding vol-cano locations and products comes from the online AIST database (AIST, 2008). This information is supplemented by observations from Ocean Drill-ing Program (ODP) cores on ash-layer frequency (Cambray and Cadet, 1994).
Geochemistry of Volcanic Rocks
In order to examine the relationship between
tectonics and magmatism in Kyushu, we inves-tigated the temporal and spatial variations in geochemical characteristics of volcanic prod-ucts by compiling reported data. The details are provided in the GSA Data Repository Fig-ures DR1–DR21.
1Key geochemical signatures can be used to
infer the source conditions for magma gen-esis beneath volcanic regions, through well- established global correlations between tectonic processes and magma chemistry (e.g., Pearce and Cann, 1971, 1973; Pearce and Norry, 1979; Pearce and Peate, 1995). Here we plot major and trace element data of Kyushu volcanic rocks on diagnostic diagrams (SiO
2 versus
alkalis, alkalis-FeO*-MgO (AFM), SiO2 versus
K2O, MnO-TiO2-P2O5, Zr-Nb-Y , Y versus Sr/Y
diagrams, and normal mid-oceanic-ridge basalt
(N-MORB)–normalized trace element pattern;
see Figures DR1–DR21 [footnote 1]). These key geochemical diagrams for Kyushu volca-nic rocks were used to classify the rock types and constrain the magmatic source and the tec-tonic background of magmatism. In particular, we focus on the following fi ve geochemically
distinctive rock types observed in Kyushu, which indicate tectonic setting as refl ected
in the nomenclature: (1) ocean island basalt (OIB); (2) island arc basalt (IAB); (3) interme-diate type between OIB and IAB (OIB/IAB); (4) high-magnesian andesite; and (5) adakitic andesite, dacite, and rhyolite, hereafter referred to as adakite.
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Mahony et al.
2204 Geological Society of America Bulletin, November/December 2011
TABLE 1. VOLCANIC ACTIVITY FROM 15 TO 10 MA
Region Date*
(Ma)Location Volcanism type Composition Geochemistry Reference
(if not AIST, 2008)Additional notes
SW 17–12 Southern Kyushu Plutons I-type and
S-type felsicN.D. Ishikawa and Nakamura, 1994; Kimura et al.,
2005Associated with mid-Miocene subduction
NW 15, 9–6 Hirado Island N.D. N.D. IAB to OIB type Uto et al., 2004 Due to an asthenospheric heat source
off NW Kyushu (Shinjo et al., 2000)
NW 15–4.3, 4.3–3.5,
3.5–0.6Iki Island Monogenetic volcanoes, mainly
scoria cones and lava fl owsAlkalic basalt OIB Sano, 1995; Uto et al., 2004
NW 14 Shimoshima N.D. N.D. HMA Nagao et al., 1992
SI and NE 14 Setouchi N.D. N.D. HMA (High Boron) Tatsumi, 2006
Note: Figure 5 shows the locations of the volcanic centers in this table. N.D.—no details were used in this study; IAB—island arc ba salt; SI—Shikoku Island; HMA—high-magnesian andesite.
*Approximate age of volcanic activity.
TABLE 2. VOLCANIC ACTIVITY FROM 10 TO 6 MA
Region Date*
(Ma)Location Volcanism type Composition Geochemistry Reference
(if not AIST, 2008)Additional notes
NW Episodes from
10 to 6Kitamatsuura Lava plateaus Tholeiitic basalt OIB to OIB/IAB NUMO, 2007 8, 100-m-thick lava plateaus. Temporal change from OIB/
IAB to OIB, Kakubuchi et al., 1994
NW 15–4.3, 4.3–3.5,
3.5–0.6Iki Island Monogenetic volcanoes, mainly
scoria cones and lava fl owsAlkalic basalt OIB Sano, 1995; Uto et al., 2004
NW 9–6 Hirado Island N.D. N.D. IAB to OIB type Uto et al., 2004
NW 7–2 Between Kitamatsuura
and ShimoshimaN.D. N.D. Low-boron HMA Shiraki et al., 2000
NW 7 Shimoshima N.D. Basalt OIB Nagao et al., 1992 7 Myr previously HMA erupted at Shimoshima
C 8 Y abakei N.D. N.D. OIB/IAB Kakubuchi and Matsumoto, 1990
SW 6.4–5.9 Nansatsu N.D. N.D. IAB Hedenquist et al., 1994 Likely marked the onset of subduction-related volcanism
Note: Figure 6 shows the locations of the volcanic centers in this table. OIB/IAB—intermediate type between OIB and island arc basal t; N.D.—no details were used in this study; IAB—island arc basalt; HMA—
high-magnesian andesite; OIB—ocean island basalt.
*Approximate age of volcanic activity.
TABLE 3. VOLCANIC ACTIVITY FROM 6 TO 2 MA
Region Date*
(Ma)Location Volcanism type Composition Geochemistry Reference
(if not AIST, 2008)Additional notes
NW 4.5–
2.49Fukuoka Monogenetic volcanoes Basalt OIB Hoang and Uto, 2003 Scattered centers
NW 4.3–3.5 Iki Island Lavas Alkalic basalt OIB Sano, 1995; Uto et al., 2004
NW 3 Higashimatsuura Lava plateau Basalt N.D. NUMO, 2007 Voluminous
NW 10–6 Kitamatsuura Lava plateaus Basalt OIB to OIB/IAB NUMO, 2007 8, 100-m-thick lava plateaus. Temporal
change from OIB/IAB to OIB
(Kakubuchi et al., 1994)
NW 2.7–2.2 Arita Stratovolcano Rhyolite N.D. NUMO, 2007
NW N.D. Between
Kitamatsuura and
UnzenN.D. N.D. IAB/ADK/Low-boron HMA N.D. Small scale
NW 2.5–0.5 Pre-Unzen Monogenetic volcanoes Basalt OIB N.D.
NW 4 Southern pre-Unzen N.D. N.D. OIB N.D.
SW 6–1.2 Hisatsu Initially Island arc andesite tuff breccias
erupted into a tectonic depression.
Plateau lavas (2.5–2 Ma)N.D. IAB/HMA Nagao et al., 1999 2.5–2 Ma “fl ood andesites”
(Nagao et al., 1995)
SW 4.8–1.6 Hokusatsu Lavas and pyroclastic cones Andesite and dacite IAB Uto and Uchiumi, 1997 Island arc-type
SW 2.5–1 Sendai Lavas and pyroclastic cones Basalt N.D. N.D.
SW 6.4–5.9 Nansatsu N.D. N.D. IAB Hedenquist et al., 1994
C From 6 Hohi Volcanic Zone Lava plateaus Andesite Some HMA Matsumoto et al., 1992; Shiraki
et al., 2000; Miyoshi et al., 2008b
C 2.8–2.1 Pre-Kuju Lava plateau Andesite N.D. N.D.
C 3.8–2.2 Pre-Aso caldera N.D. N.D. HMA (3.8 Ma) IAB (2.2 Ma) Miyoshi, 2008 Increased K2O, B/Nb and other fl uid-mobile
elements with time
Note: Figure 8 shows the locations of the volcanic centers in this table. N.D.—no details were used in this study; OIB/IAB—intermedia te type between OIB and island arc basalt; IAB—island arc basalt; ADK—
adakite; HMA—high-magnesian andesite; OIB—ocean island basalt.
*Approximate age of volcanic activity.
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2205
TABLE 4. VOLCANIC ACTIVITY FROM 2 TO 0 MA
Region Date*
(Ma)Location Volcanism type Composition Geochemistry Reference
(if not AIST, 2008)Additional notes
NW 1.7–1.4, 1–0.6 Iki Island Monogenetic volcanoes,
scoria conesAlkalic basalt OIB Sano, 1995; Uto et al., 2004
NW 1.07–0.66 Goto Islands
(Ojikajima,
Akadaki,
Kishuku, Fukue,
Kyonotake,
Onidake,
Hinodake)Monogenetic volcanoes,
lava fl ows, cinder cones,
pyroclastic conesBasalt OIB N.D.
NW From 1.3 Ukujima Polygenetic stratovolcano Basalt, andesite, rhyolite OIB Sudo et al., 1998
NW 1.1 Kurose Volcanic neck (remains) Basalt OIB N.D.NW 1.3–1, 0.8–0.4 Taradake Stratovolcano, lava domes Andesite, basalt, dacite OIB/ OIB/IAB N.D.NW Formed 0.5–0.3 Unzen Lava fl ows and domes
(0.5–0.3 Ma)N.D. OIB/IAB Sugimoto et al., 2005
C 0.35–0.02 Himeshima Lava domes and pyroclastic
fl o w sDacite, rhyolite ADK Itoh, 1990
C 1.5–1.1 Futago Lava domes and lava fl ows Andesite, dacite ADK Nakada, 1986
C 1.1 Kanagoe Stratovolcano Andesite N.D. N.D.C 0.4–0.3, 0.015 Hiji Lava fl ows Andesite, dacite (0.4–0.3) IAB Kamata, 1987, 1989
C 0.04 Yufu Lava dome stratovolcanoes Andesite, dacite OIB/IAB /
ADK/IABKobayashi, 1984; Sugimoto et
al., 2006
C Approx. 2 Pre-Tsurumi Lava plateau Andesite N.D. N.D.
0.04 Tsurumi Lava dome stratovolcanoes Andesite, dacite OIB/IAB /
ADK/IABKobayashi, 1984; Sugimoto et
al., 2006
C N.D. FT/TM/OA/NH N.D. N.D. N.D. N.D.
C 1–0.7, 0.7–0.3 Shishimuta Caldera 1–0.7, lava plateau; 0.7,
caldera forming; 0.7–0.3,
small volume lavasN.D. Kamata et al., 1988; Kamata,
1989Marked change at 0.7 Ma from N-S extension to N-S
compression, coincides with change in volcanism style
C 0.6–0.4 Waita-yama Stratovolcano, lava domes Andesite N.D. Kamata, 1998
C N.D. Hane-yama N.D. N.D. N.D. Daishi, 2006
C N.D. Pre-Kuju N.D. Rhyolite N.D. Daishi, 2006
0.6 Kuju Andesite volcano complex N.D. ADK/IAB N.D.
C 1.6, 0.7 Hohi volcanic
zone1.6 Ma K
2O increase in lavas. N.D. N.D. Kamata et al., 1988; Kamata,
1989
C 0.45 Pre-Aso caldera Andesite pyroclastic fl ows,
pyroclastic cones, lava fl ows,
biotite rhyolitesN.D. ADK/IAB Kaneoka and Ozima, 1970
C 0.3, Aso 1; 0.14, Aso
2; 0.12, Aso 3; 0.09,
Aso 4Aso Caldera-forming eruptions,
pyroclastic fl ow depositsRhyolite, dacite N.D. Ono and Watanabe, 1985;
Hunter, 1998
From 0.022 Naka-dake,
Take-dake (Aso)Postcaldera central
stratoconesBasalt, andesite N.D. Miyabuchi et al., 2004 Post–Aso caldera volcanism
C 0.09 Omine Intracaldera (Aso) volcanism
formed fl ank monogenetic
volcanoDacite N.D. Matsumoto et al., 1991 Active directly before Aso 4, formed along a fault line
trending SW of Aso
C 0.15 Akai Intracaldera (Aso) volcanism
formed fl ank volcanoDacite N.D. Matsumoto et al., 1991 Active between Aso 1 and Aso 2, formed along a fault line
trending SW of Aso
C 1.1 Kimpo Stratovolcano Andesite OIB/IAB N.D.
C 0.022–0.006 Oninomi N.D. N.D. OIB/IAB Ohta et al., 1992 Small volume
SW 6–1.2 Hisatsu N.D. N.D. IAB Nagao et al., 1999
SW 3–1, episodic activity
until 0.025Hokusatsu/pre–
Aira calderaLava fl ows Andesite, basalt IAB/OIB(1
Ma)Uto and Uchiumi, 1997; Sudo
et al., 2000
SW 0.025 Aira Caldera-forming eruption Rhyolite N.D. Aramaki, 1984; Tsukui and
Aramaki, 1990Produced the 110 km3 rhyolitic Ito pyroclastic fl ow deposit
SW 0.022–0 Sakurajima Postcaldera stratovolcano Andesite, dacite N.D. N.D. Formed on the southern rim of Aira caldera
SW 0.016 Wakamiko Caldera Rhyolite N.D. Shimomura, 1960 6 × 3 km submarine caldera in the north eastern corner of
the Aira caldera
SW 0.5–0.35 Imuta Lava dome Dacite N.D. N.D.
SW 0.5 Satsuma-
MaruyamaTwo lava domes Andesite N.D. N.D.
SW 0.1 Aojiki Maar N.D. N.D. N.D.
SW 0.007 Y onemaru-
Sumiyoshi-ikeMaar Basalt N.D. N.D.
SW 1.6, 1.4 Pre-Kirishima Lava shield Andesite N.D. N.D.
(Continued )
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2206 Geological Society of America Bulletin, November/December 2011
TABLE 4. VOLCANIC ACTIVITY FROM 2 TO 0 MA ( Continued )
Region Date*
(Ma)Location Volcanism type Composition Geochemistry Reference
(if not AIST, 2008)Additional notes
SW 0.3–0.15, 0.1 Kirishima Shield volcano (0.3–0.15 Ma),
postcaldera stratovolcano
since 0.1 MaAndesite IAB N.D.
SW 0.5–0.4 Kobayashi Caldera Dacite, rhyolite N.D. N.D. Associated with Kirishima volcano
SW 0.3 Kakuto Caldera Dacite, rhyolite N.D. N.D. Associated with Kirishima volcano
SW 0.11 Ata Caldera N.D. N.D. Matsumoto and Ui, 1997 25 km (E-W) ×12 km submarine caldera in the mouth of
Kagoshima bay
SW 0.006 Ikeda-ko Caldera, lava dome, maar Rhyolite, dacite, andesite N.D. N.D. Lies directly west of Ata calderaSW 0.004 Kaimondake Stratovolcano, lava dome Basalt, andesite IAB N.D. Lies directly west of Ata caldera
SW 0.095, 0.0063 Kikai Caldera-forming eruptions
producing large ignimbritesN.D. N.D. Ono et al., 1982; Walker et
al., 1984; Machida, 199920 × 17 km diameter caldera, 40 km SW of Kyushu.
The 0.006 Ma VEI 7 eruption was the largest Holocene
eruption in Japan
SW 0.7 Y ahazu-dake,
TakeshimaPre–Kikai caldera volcanism N.D. N.D. Newhall and Self, 1982;
Torrence and Grattan, 2002Located on the Kikai caldera rim
SW 1–0.9 Kuroshima Stratovolcano Andesite N.D. N.D.
SW 0.7 Tairajima,
SuwanosejimaStratovolcano Andesite N.D. N.D.
SW 0.01–0 Approx. 16 new
centersLava domes,
Stratovolcanoes, small
calderasAndesite, rhyolite N.D. N.D.
Note: Figure 9 shows the locations of the volcanic centers in this table. N.D.—no details were used in this study; OIB/IAB—intermedia te type between OIB and island arc basalt; ADK—adakite; IAB—island arc
basalt. Volcanoes FT/TM/OA/NH: FT—Fukuman-yama and Tateishi-yama; TM—Takahira-yama and Mizuguchi-yama; OA—Ojika-yama and Amagoi -dake; NH—Noine-dake and Hanamure-yama.
*Approximate age of volcanic activity.Ocean island basalt–type rocks are character-
istic of an enriched asthenosphere source (e.g., Saal et al., 1998; Parman et al., 2005; Jackson et al., 2008). These rocks are plotted on the fi elds of ocean island tholeiite and/or ocean
island alkaline in MnO-TiO
2-P2O5 diagram, and
are plotted on the fi elds of within-plate tholei-
ite and/or within-plate alkaline, and/or plume (P) -MORB fi elds in Zr-Nb-Y diagram.
Island arc basalts are enriched in fl uid mobile
large ion lithophile elements (e.g., K, Rb, and Ba) and light rare earth elements, and are depleted in fl uid immobile high fi eld strength
elements (e.g., Nb, Zr) and heavy rare earth ele-ments. Island arc basalts are plotted on the fi elds
of island arc tholeiite and calc-alkaline basalt in MnO-TiO
2-P2O5 diagram, and are plotted on the
fi eld of volcanic arc basalt in Zr-Nb-Y diagram.
We defi ned OIB- and IAB-type as the rocks that
plot around the boundary area between OIB and island arc basalt in the MnO-TiO
2-P2O5 and
Zr-Nb-Y diagrams, implying contributions from both OIB and island arc basalt sources. We also describe rocks as OIB/IAB if they are defi ned as
OIB in one diagram and as island arc basalt in another diagram.
Boron (B) is a useful element for identify-
ing subducted oceanic-slab infl uence on the
subarc mantle compositions, because B is enriched in altered oceanic crust and sea-fl oor
sediment (e.g., Ishikawa and Nakamura, 1993; Smith et al., 1995), and is selectively parti-tioned into fl uid phase during fl uid-fl ux melt-
ing at the base of mantle wedge (e.g., Moran et al., 1992; Bebout et al., 1999). In contrast, OIB and MORB have low B contents (Sun and McDonough, 1989; Ryan and Langmuir, 1993; Chaussidon and Jambon, 1994; Chaussidon and Marty, 1995; Ryan et al., 1996; Leeman and Sis-son, 1996) refl ecting an absence of a subduct-
ing slab component. B/Nb (Fig. 4) is a sensi-tive indicator of slab involvement. Boron and Nb have similar solid and/or melt distribution coeffi cients so that B/Nb are not signifi cantly
affected by partial melting and crystal fraction-ation. However, B and Nb have entirely differ-ent chemical behavior in fl uid-related processes
with B having a much higher mobility than Nb. Boron/Nb variation also does not refl ect crustal
assimilation, because the ratio in crust is vanish-ingly low (Ishikawa and Nakamura, 1994; Lee-man and Sisson, 1996; Ishikawa and Tera, 1997; Ishikawa et al., 2001; Sano et al., 2001; Tonarini et al., 2004).
Here we classify rocks containing more
than 53 wt% of SiO
2 and 6–8 wt% of MgO as
high-magnesian andesite by following Shiraki (1993). Setouchi high-magnesian andesite mag-mas, distributed in southwest Japan, are inter-preted as either due to hydrous partial melting
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2207of the mantle or due to the interaction between
the mantle and silicic slab melt in subduc-tion zones (e.g., Tatsumi, 2003). The origin of Setouchi high-magnesian andesite magmas are often ascribed to the subduction of a young and hot oceanic plate (Shikoku Basin) and backarc opening (Tatsumi, 2003, 2006). However, the origin of the Kyushu high-magnesian andesite is still controversial. Previous work suggests that Kyushu high-magnesian andesite was generated independently from subduction, by interaction between injected asthenospheric mantle and preexisting metasomatized lithospheric mantle, on the basis of their mature stage of volcanism and coexistence of high-magnesian andesites
and OIB-type rocks (Kakubuchi et al., 1995; Shiraki et al., 2000; Miyoshi et al., 2008). We contend that the exact origin and plate tectonic context of both the Setouchi and Kyushu high-magnesian andesites are not well understood.
Adakitic silicic rocks are characterized by
high Sr/Y ratios (see Fig. 4). Hypotheses for the generation of adakites remain controversial (e.g., Castillo, 2006). Defant and Drummond (1990) suggested that adakites are formed by direct melting of the young (<25 Myr) hot sub-ducting slab or sediments. Alternatively adakite magma could be generated by the partial melt-ing of the mafi c lower crust under garnet and/or
amphibole stable conditions (e.g., Atherton and Petford, 1993). In the case of Kyushu, high Sr/Y adakites were distributed in the volcanic front and backarc area. The volcanic front adakites in Yufu at 0.04 Ma were interpreted by Sugi-moto et al. (2006) to be generated by mixing between MORB mantle and partial melt of the subducted terrigenous sediments on the basis of calculation using trace elements and isotopic data. Shiraki et al. (2000) argued the Pliocene backarc adakites are generated by interaction between injected asthenospheric mantle and preexisting metasomatized lithospheric mantle,
128° 129° 130° 131° 132° 133° 134° 135°31°32°33°34°35°
−5000 072 mm/yr
AMUR/PSP Nankai Trough
PHILIPPINE SEA
PLATEMTL
Okinawa TroughAMURIANPLATE
KyushuSouthwest
Honshu
Shikoku Island
KGBSGFG
KJ
AsoHM
AR
U
KS
AI
IK
AT
KKShimabara
PeninsulaBeppu
Kagoshima BayNagasakiGoto
IslandsFukuoka
Kumamoto
SJOita
OKTLNon-
Volcanic
Region
Southern Volcanic RegionNorthwestern
Volcanic Region
HVZ
−5000 0
metersTsushima
Islands
CentralVolcanic Region
N
0 50 kmForearcBack-arc
Figure 3. Detailed active tectonic setting of Kyushu. Several key cities, volcanoes, and tectonic features are highlighted on t his
map. Black text indicates a city or other geographical location. Red triangles mark key volcano locations. Tectonic features ha ve
blue text; the two grabens are bounded by dotted blue lines. Green text indicates the three main volcanic regions and the non-volcanic region. The forearc region is defi ned as the area between the volcanic front and the trench, the volcanic front marked
here by the volcanoes FG-KJ-Aso-KS-AI-SJ-IK-AT-KK. The backarc region is that behind the volcanic front. Abbreviations: FG—Futago volcano; HVZ—Hohi volcanic zone, marked by red line; HM—Higashimatsuura volcanics; AR—Arita volcanic rocks; U—Unzen volcano; KJ—Kuju volcano; Aso—Aso volcano; KS—Kirishima volcano; AI—Aira caldera; SJ—Sakurajima volcano; IK—Ikeda-ko volcano; AT—Ata caldera; KK—Kikai caldera; BSG—Beppu-Shimabara graben; MTL—Median Tec-tonic Line; OKTL—Oita-Kumamoto tectonic line; KG—Kagoshima graben.
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Mahony et al.
2208 Geological Society of America Bulletin, November/December 2011024 6 0.1 0.2 0.3 20 40 60
B/Nb B/Zr Sr/YOninomi
Yufu
Kuju
Aso
Kirishima
Mantle
values 31°32°33°34° LatitudeFutagoHimeshima
Aira
Imuta
Figure 4. Along-arc variations of B/Nb, B/Zr, and Sr/Y ratios of volcanic products from
around the recent volcanic front in Kyushu. Black circles in B/Nb and B/Zr diagrams show data for the basaltic rocks. White circles in Sr/Y diagram show data for the hornblende-bearing silicic products.on the basis of the coexistence of adakites and
OIB-type rocks.
Active Volcanic Regions
Here we have divided Kyushu into three
regions of active volcanism: the northwestern, central, and southern (Fig. 3). We defi ned these
regions as areas with similar volcanic history. A common major change of volcanism in these three volcanic regions is the widespread lava plateaus that preceded other eruption styles. The northwestern volcanic region is dominated by episodic backarc volcanism, with many mono-genetic volcanoes and lava plateaus (we defi ne
lava plateaus as a series of dominantly long, runout lava eruptions from distributed sources). The central volcanic region is a long-lived area of volcanism (including the recent onset of arc volcanism), changing in style through time. Types of volcanism in the central volcanic region have included voluminous lava plateaus, lava domes, cinder cones, and calderas (such as the 18 km × 25 km Aso caldera, Fig. 3). The southern volcanic region experienced wide-spread lava plateau volcanism throughout the Pliocene, but since 0.3 Ma there has been volca-nic front caldera formation and postcaldera stra-tovolcanoes (e.g., Aira caldera 18 km × 20 km, Sakurajima postcaldera stratovolcano, Fig. 3). There is a fourth region of Kyushu, the “nonvol-canic” region located partially in the northern forearc in Figure 3. This region has no record of volcanism, although there are Miocene I-type (ca. 12–15 Ma) and S-type (14–17 Ma) granitic plutons (e.g., Kimura et al., 2005). There is also a marked peak in frequency of voluminous vol-canic ash layers in ODP sites 296, 442, 443, and 445 in the Shikoku Basin (Cambray and Cadet, 1994), suggesting substantial volcanism in southwest Japan in the Miocene.
Volcanic History 15–10 Ma
The I- and S-type plutons (Fig. 5) form part
of more widespread igneous activity (Table 1) in southwest Japan between 17 and 13 Ma (Kano et al., 1991; Kimura et al., 2005). Furukawa and Tatsumi (1999) proposed that the Setou-chi high-magnesian andesite volcanism formed due to subduction of the young, hot Philippine Sea plate between 17 and 10 Ma (Takahashi, 1981; Tatsumi, 1983). Tatsumi (2006; Tatsumi et al., 2001; Tatsumi, 1982) suggested that the Setouchi high-magnesian andesite volcanism is associated with the opening of Japan Sea and subduction of hot Shikoku Basin lithosphere. High-magnesian andesite volcanism also occurred in Kyushu almost simultaneously with the Setouchi high-magnesian andesites at Shi-moshima (Nagao et al., 1992; Fig. 5).
Iki Island
15–4.3 Ma
Hirado15 Ma, 9–6 Ma
Possible location of heat source, 11 MaSetouchi 14 Ma
Shimoshima14 MaFelsic plutons
N
0 50 km130° 132°
34°
31°II
I
IS
S
Figure 5. Volcanism in Kyushu from 15 to 10 Ma. Shaded areas of varying size represent
lava plateaus, or in some cases represent volcanism of an unknown nature. Triangles rep-resent stratovolcanoes, or monogenetic volcanoes. There is a band of I- and S-type felsic plutons (represented by black dots with an I and an S, respectively) trending northeast-southwest across central Kyushu. There is minor volcanism in the northwestern and cen-tral volcanic regions. The locations of the I- and S-type felsic plutons are after Kimura et al. (2005).
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2209Volcanic History 10–6 Ma
The 10–6 Ma period (Fig. 6; Table 2) is
marked by a lack of obvious subduction-related volcanism (Fig. 7), which has led to the sug-gestion that there was a hiatus in subduction (Kamata and Kodama, 1994). Low rates of explosive volcanism are indicated by the low abundance of volcanic ash layers in ODP cores from the Shikoku Basin (Cambray and Cadet, 1994). Effusive volcanism mostly occurred in the northwestern region, with minor late volca-nism in the southern region (Fig. 6).
Volcanic History 6–2 Ma
From 6 to 2 Ma, volcanic activity in Kyushu
increased relative to the 10–6 Ma period (Fig. 8; Table 3), with notable formation of many andes-itic lava plateaus in areas of future caldera volca-nism. The northwest volcanic region continued forming basaltic lava plateaus and monogenetic volcanoes, as well as new rhyolitic volcanism. V oluminous andesitic lavas characterized the
early stage of the Hohi volcanic zone (6–2 Ma) (Kamata, 1989). The early stages erupted high-magnesian andesite at the Hohi volcanic zone tectonic edges, followed later (2–0 Ma) by island arc basalts (Kamata, 1987; Nakada and Kamata, 1991), which focused toward the center of the Hohi volcanic zone (Kamata and Kodama, 1994). These lavas became compo-sitionally more K enriched with time and after peaking in activity at 5 Ma, the Hohi volcanic zone decreased in activity throughout the Plio-cene to the present (Kamata and Kodama, 1994). Around the same time, volcanism increased in the Shimabara peninsular area (Yokose et al., 1999), the eventual location of Unzen volcano. On the western edge of the Hohi volcanic zone, Aso volcano erupted high-magnesian andesites at 3.8 Ma and island arc basalts at 2.2 Ma, which showed an increase in K
2O, B/Nb, and other
fl uid-mobile elements content with time (Fig. 4; Miyoshi, 2008). High-magnesian andesites have
low B/Nb ratios that are similar to MORB or OIB (0.05–0.5; Ryan et al., 1996). On the other hand, the island arc basalts show signifi cantly
higher B/Nb than high-magnesian andesites. Therefore the elevated B/Nb suggests that the addition of slab-derived fl uid to the source man-
tle beneath the Aso area occurred between 3.8 and 2.2 Ma (Miyoshi, 2008). In this western part of the central volcanic region, the adakite and island arc basalt magmas coexisted with high-magnesian andesites. Pre-caldera volcanism in Aso was dominated by andesite lava plateaus (Ono and Watanabe, 1985).
In the southern volcanic region, andesite lava
plateaus developed in the Nansatsu, Hokusatsu, and Hisatsu areas, parallel to and westward of the Kagoshima graben (Nagao et al., 1999) (Fig. 8; Table 3). Ocean Drilling Program sites 296, 442, and 443 adjacent to Kyushu in the Shi-koku Basin show a high frequency of volcanic ash layers from 6 Ma to the present (Cambray and Cadet, 1994).
Volcanic History 2–0 Ma
The most recent episode of arc volcanism in
Kyushu started at the beginning of the Quater-nary (e.g., Watanabe, 2005). Signifi cant changes
in the mode of eruption and whole-rock chem-istry occurred throughout Kyushu (Kamata and Kodama, 1994; see Fig. 7 for a summary). An active volcanic arc related to Philippine Sea plate subduction formed in the central and southern regions of Kyushu, extending from Aso and trending northeast-southwest (Fig. 3). Arc volcanism has been prevalent in Kyushu for the past 2 Myr, with a marked pulse of large-magnitude explosive eruptions with associated caldera formation from 0.3 Ma (e.g., Aso, Aira, Kakuto, Ata, and Kikai; Fig. 9). The large calde-ras in the arc were mostly preceded by Pliocene to early Pleistocene voluminous andesitic lava fl ows. Ocean Drilling Program data on Shikoku
Basin cores also show a marked increase in ash layer frequency after 2 Ma, especially in site 442 (Cambray and Cadet, 1994).
The northwestern volcanic region showed
continued backarc volcanism, typically with new monogenetic volcano centers as well as Unzen volcano, which lies in a transitional zone between the arc and backarc regions. There is a nonvolcanic gap in the volcanic front to the southeast of Unzen. Taken together, the location of Unzen and the gap are enigmatic. The central volcanic region was the location of the fi rst Quaternary caldera volcanism in
Kyushu with Shishimuta caldera, followed by multiple caldera-forming eruptions at Aso volcano (Fig. 9; Table 4). The Hohi volcanic zone saw a marked increase in K
2O content in Kitamatsuura
basalts, 10–6 Ma Iki Island
Shimoshima
basalts, 7 Ma
Nansatsu
6.4–5.9 MaYa b a ke i8 Ma
Hirado9–6 Ma
Volcanic activity7–2 Ma
N
0 50 km130° 132°
34°
31°
Figure 6. Volcanism in Kyushu from 10 to 6 Ma. Shaded areas of varying size represent
lava plateaus, or in some cases represent volcanism of an unknown nature. Triangles represent stratovolcanoes, or monogenetic volcanoes. Dashed gray lines in the central region mark the Beppu-Shimabara graben. Episodes of effusive volcanism occurred sporadically through time, mainly located along the very western edge of Kyushu (i.e., Nansatsu, Kitamatsuura). The location of the Nansatsu volcanic rocks is taken from Shinjo et al. (2000).
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Mahony et al.
2210 Geological Society of America Bulletin, November/December 20110 100 km130°E10–6 Ma
32°N33°N
6–2 Ma
130ºE130ºE130°E2–0 Ma
32°N33°NIki
Aso
AsoKuju
UnzenUnzenHimeshimaHimeshimaY abakei
ShimoshimaShimoshima Shimoshima
Nansatsu
FukuokaIki
Kurose
Ojikajima
FukueLeading Leading
edge?edge?
Yufu-TsurumiHimeshima
Futago130°E15–10 Ma
32°NHirado
OIB
OIB/IABIABHMA
ADK
15–7 Ma
Felsic plutons17–12 Ma
130°E33°N
32°N10–6 Ma
8 Ma
7 Ma
<4 MaShimoshima
14 MaLeading Leading
edge?edge?14 Ma
7–2 Ma
Low-Boron HMALow-Boron HMA Low-Boron HMASetouchi HMASetouchi HMA
(High-Boron)(High-Boron)
Setouchi HMA(High-Boron)
5–3 Ma
2.2 Ma
3.8 MaLow-Boron HMALow-Boron HMA Low-Boron HMA
4 Ma4 Ma4 MaLow-Boron HMALow-Boron HMA Low-Boron HMA
1 Ma1 Ma1 MaHisatsu4–1 MaIIIIKitamatsuura
NSS
KirishimaKirishima KirishimaTaradakeTaradake TaradakeTaradake Taradake
UnzenUnzen Unzen
HokusatsuHokusatsuHokusatsu
KaimonKaimon KaimonHokusatsuHokusatsu
NansatsuNansatsuHisatsuHisatsu
Hokusatsu
NansatsuHisatsu
Figure 7. Concise summary of the geochemistry of volcanoes in Kyushu through the past 15 Ma. The dots represent
analyzed samples from named volcanoes and/or areas of volcanism that have been classifi ed by type (different shad-
ing). Solid black dots are ocean island basalt (OIB); dark gray dots with white spots are hybrids between OIB and island arc basalt (IAB). Plain gray dots are IAB; pale-gray dots with dark-gray spots are high-magnesian andesite (HMA); and white with dark-edge dots are adakites (ADK). Black circles with an I or S in them, respectively, rep-resent the locations of I- and S-type felsic plutons (from Kimura et al., 2005). Where the color is not in a small circle (e.g., large Fukuoka area on the 6–2 Ma map), the volcanism is distributed throughout this area, but not necessarily spatially continuous.
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2211lavas by 1.6 Ma, which preceded the marked
change from north-south extension to north-south compression in the Hohi volcanic zone at 0.7 Ma (Kamata et al., 1988), coinciding with the change in volcanism style from cal-dera formation (Shishimuta caldera, 8 km wide, >3 km deep; Kamata, 1989) to small-volume lavas and domes. The southern volca-nic region has seen an eastward migration of the volcanic front since 6 Ma. The Nansatsu, Hokusatsu, and Hisatsu volcanic areas, char-acterized by island arc basalts, represent the location of the early volcanic front (Watanabe et al., 1992). The Kagoshima graben represents the current location (Figs. 7 and 9). Southern Kyushu has had several Quaternary caldera-forming eruptions, at Kakuto, Kobayashi, Aira, Wakamiko, and Ata calderas. Kirishima vol-cano is a postcaldera volcano associated with
Kakuto and Kobayashi calderas, comprising a series of cones on its fl anks (International V ol-
canological Association, 1962), all formed in an elliptical volcano zone (~30 km × 20 km) trending northwest-southeast. Interestingly, the location and northwest-southeast orienta-tion of Kirishima and its associated vents coin-cide and align with the left-lateral shear zone identifi ed from GPS observations (Wallace et
al., 2009a) that crosscuts southern Kyushu. Farther south, offshore Kyushu in the Ryukyu arc (Fig. 9), the Kikai caldera had multiple caldera- forming eruptions, including a VEI 7 eruption at 0.063 Ma (Ono et al., 1982; Walker et al., 1984; Machida, 1999). Many volcanic centers in the Ryukyu arc became active dur-ing the Quaternary (Fig. 9; Table 4).
Kitamatsuura
basalts Iki Island 4.3–3.5 Ma
Higashimatsuura basalts, 3 Ma
Arita rhyolites 2.7–2.2 MaHVZ formation
6 Ma HVZ andesite lava plateau
Okinawa
Troughopening Aso, 3.8–2.2 Ma
Nansatsu
6.4–5.9 MaHokusatsu 4.8–1.6 MaHisatsu 6–1.2 MaFukuoka 4.5–2.49 Ma
Volcanic activity, 4 MaVolcanic activityPre-Kuju2.8–2.1 Ma
Pre-Unzen
2.5 Ma
Sendai
2.5–1 MaN
0 50 km130° 132°
34°
31°
Figure 8. Volcanism in Kyushu from 6 to 2 Ma. Shaded areas of varying size represent lava
plateaus, or in some cases represent volcanism of an unknown nature. Triangles represent stratovolcanoes, or monogenetic volcanoes. Reinitiation of subduction at ca. 6 Ma is closely followed by the formation of the Hohi volcanic zone (HVZ; marked by dotted gray line). Volcanism in the HVZ initially formed at the edges of the zone and focused into the center with time. Extensive andesite lava plateaus also formed in the HVZ. Widespread effusive volcanism continued in the northwestern and southern regions. Dashed gray lines in the central region mark the Beppu-Shimabara graben; dashed and solid lines in the southern region mark the Kagoshima graben. The sporadic lavas located in the Fukuoka area are from scattered, small monogenetic volcanoes. The exact locations of these volcanoes are unknown; so the locations of the dated samples are plotted here instead of volcano locations. The locations of the Hisatsu, Hokusatsu, and Nansatsu volcanic rocks are taken from Shinjo et al. (2000).TECTONIC CONFIGURATION OF
THE KYUSHU REGION FROM 15 MA TO PRESENT
In order to understand the changing char-
acter of volcanism in Kyushu, the evolution of the plate-boundary confi guration must be
reconstructed. Philippine Sea plate motion has profoundly impacted the tectonic and volcanic evolution of southwest Japan. According to paleomagnetic data, the Philippine Sea plate has undergone large northward translation and clockwise rotation (~90° relative to the sur-rounding plates) since 50 Ma (Hall et al., 1995a, 1995b, and references therein), although there is controversy regarding the exact timing of the northward movement of the Philippine Sea plate (Yamazaki et al., 2010). The Philippine Sea plate has also undergone notable changes in its kinematics, likely due to major regional tec-tonic events, including collisions on its bound-aries (Eguchi, 1984; Hall et al., 1995a, 1995b; Sdrolias et al., 2004). Hall et al. (1995a, 1995b) suggested that between 40 and 50 Ma the Philip-pine Sea plate underwent ~50° clockwise rota-tion, negligible rotation from 25 to 40 Ma, ~34° clockwise rotation from 5 to 25 Ma, and ~5.5° clockwise rotation from 5 Ma to present. Cur-rently, geodetic data indicate ~1°/Myr clockwise rotation rate for the Philippine Sea plate (Sella et al., 2002), generally consistent with the 5 Ma to present paleomagnetic estimates of vertical axis rotation rate. At 2 Ma, a shift in Philippine Sea plate motion occurred, causing Philippine Sea plate/Amurian plate relative plate motion to become more oblique in southwest Japan (in a dextral sense), leading to Pliocene reactivation of the Median Tectonic Line as a right-lateral strike-slip fault (Kamata and Kodama, 1994; Itoh et al., 1998; Kamata, 1998). Studies of Philippine Sea plate kinematics suggest that a major northward shift in the position of the Phil-ippine Sea plate/Eurasia relative pole of rotation from 15°N to 48°N occurred at ca. 5 Ma (e.g., Hall et al., 1995a), which would have caused a major change in relative motion at the southwest Japan–Philippine Sea plate boundary around 5 Ma.
Alternative models for Philippine Sea plate
kinematics have recently been developed from the interpretation of paleomagnetic data from sites on the northern portion of the Philippine Sea plate (Yamazaki et al., 2010). These new data are interpreted to suggest that the Philip-pine Sea plate underwent most of its north-ward movement prior to 15 Ma (suggesting that between 50 and 15 Ma, the Philippine Sea plate rotated 90° clockwise about a pole near 23°N/162°E, and that northward movement of the Philippine Sea plate after 15 Ma is negligible
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Mahony et al.
2212 Geological Society of America Bulletin, November/December 2011Iki island
Kishuku (Fukue) Ukujima
Tara
-dakeKurose Futago Kanagoe
Hokusatsu Hisatsu
Okinawa TroughKyonotakeOjikajima FT
TM
OA Hiji
Kuroshima
Tairajima Imuta
Satsuma-Maruyama Satsuma-Maruyama Satsuma-Maruyama Kobayashi CAShishimuta CAKimpo
NH
Waita-yama Hane-yama
Akadaki
Onidake AkaiOmineHimeshima
Aojiki
Ata CA Aira CA Wakamiko CA Kakuto CAHinodake Aso KujuYufu
Tsurumi
Unzen
Kirishima
Sakurajima
Kaimondake
Satsuma-Io JimaSatsuma-Io Jima
Satsuma-Io Jima
Satsuma-Io Jima
Kuchinoerabu-jimaKuchinoerabu-jima Kuchinoerabu-jimaKuchinoshima
Nakano-shima
Suwanose-jima Oki and
Suwanose-jima
Akesuki-JimaKogaja-Jima andKogaja-Jima and
Gaja-JimaGaja-JimaKogaja-Jima and
Gaja-JimaKikai CA Ikeda-ko CA Y onemaru-Sumiyoshi-ikeYonemaru-Sumiyoshi-ike Y onemaru-Sumiyoshi-ike
2–1 Ma
1–0.3 Ma
0.3–0.01 Ma
0.001–0 Ma
N
0 65 km130° 132°
34°
31°
Figure 9. Volcanism in Kyushu from 2 to 0 Ma. Shaded areas of varying size represent lava plateaus, or in some
cases represent volcanism of an unknown nature. Triangles represent stratovolcanoes, or monogenetic volca-noes; stars represent calderas. The Quaternary is shaded into four sections: 2–1 Ma (white with dark outline), 1–0.3 Ma (pale gray), 0.3–0.01 Ma (dark gray), and 0.001–0 Ma (black). Many lava shield volcanoes and lava plateaus formed at ca. 2 Ma (nontriangle or star “areas” of color here), in locations of future caldera volcanism. From 1 to 0.3 Ma, the major change in volcanism is the start of caldera formation in Kyushu, at Shishimuta and Kobayashi calderas (represented by larger stars). The Hohi volcanic zone (HVZ) gains an infl ux of stratovolcanoes
from 0.6 Ma in the Kuju volcano complex; FT—Fukuman-yama and Tateishi-yama; TM—Takahira-yama and Mizuguchi-yama; OA—Ojika-yama and Amagoi-dake; NH—Noine-dake and Hanamure-yama. From 0.3 Ma, volcanism in Kyushu is dominated by large calderas, namely Aso, Kakuto, Aira, Ata, and Kikai (calderas are represented by large stars). These calderas all form within grabens, the Beppu-Shimabara graben in the central region (outlined by dashed/dot and dashed lines) and the Kagoshima graben in the southern region (outlined by solid black lines). Volcanism in the southern volcanic region generally youngs southward. Holocene volcanism has seen the volcanic front extend down in the Ryukyu arc, with the largest Holocene eruption occurring at Kikai volcano. Through the Holocene, the long-lived centers of Fukue (Hinodake), Unzen, Aso, Kuju, Yufu/Tsurumi, Kirishima, and Sakurajima continue to have active stratovolcanoes. The Holocene has had no widespread lava plateau eruptions. The locations of the Hisatsu and Hokusatsu volcanic rocks are taken from Shinjo et al. (2000).
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2213(Yamazaki et al., 2010). However, the Yamazaki
et al. (2010) data set does not include samples between 25 and 10 Ma, so it is diffi cult to assess
the timing of post–25 Ma changes in Philippine Sea plate motion from these data reliably. We suggest that these data (and their associated uncertainties) can be fi t equally well by 0.4°–
0.5°/Myr (in latitude) of northward movement of the Philippine Sea plate from 29 Ma to pres-ent, comparable to the ~0.7°/Myr northward movement of the Philippine Sea plate from 5 to 15 Ma that we expect in the reconstructions pre-sented here using rotation poles from previous studies of Philippine Sea plate kinematics.
Methods
To reconstruct the plate-boundary confi gura-
tion for southwest Japan over the past 15 Ma, published fi nite rotation poles for the Philip-
pine Sea plate relative to Eurasia from Gaina and Müller (2007) are used for the 2–15 Ma reconstructions. Note that the Gaina and Mül-ler (2007) Philippine Sea plate–Eurasia poles are largely derived from kinematic studies by Seno et al. (1993), Hall et al. (1995a, 1995b), and Sdrolias et al. (2004). For 2 Ma to present, poles of rotation for the Philippine Sea plate, the Amurian plate, and other tectonic plates and blocks in the region are derived from inversion of earthquake slip vectors and contemporary GPS data (Wallace et al., 2009a). Paleomag-netic data from the Kyushu forearc (Kodama and Nakayama, 1993; Kodama et al., 1995) are used to constrain the anticlockwise rotation of the Kyushu forearc from ca. 6 Ma to present. There remains considerable uncertainty about when Kyushu forearc rotation began: Kamata and Kodama (1999) proposed that it started in the Quaternary (<2.5 Ma), suggesting the rota-tion was very rapid (~15°/Myr). Other stud-ies suggest that the rotation has been ongoing since 6 Ma (Kato et al., 1998; Yamaji, 2003), as this is consistent with evidence for exten-sional deformation in southern Kyushu for the past 5 Ma, and slab rollback of the Philippine Sea plate since at least 5 Ma (Yamaji, 2003). For the purposes of this study, we assume that anticlockwise rotation of the Kyushu forearc has been ongoing since 6 Ma, at a constant rate (~5°/Myr). In our reconstructions, the positions of major physiographic features are tracked, such as the Kyushu-Palau ridge, the Izu-Bonin-Mariana arc, and the Shikoku Basin spreading center. We assume that the position of the Izu-Bonin Trench remains in the same location that it is today relative to the Izu-Bonin-Mariana arc throughout the time period of this reconstruc-tion. A similar assumption is made for the posi-tion of the Ryukyu and Nankai Troughs relative
128°E 130° 132° 134° 136° 138° 140° 142° 144°22°24°26°28°30°32°34°36°N
~ 8cm/a
~8 cm/aSubdu
ctionSubduction
Pacific
plate
Philippine
Sea
plateEurasian/Amurian plate
KPR (15 Ma)KPR (10 Ma)SBSC (15 Ma, has ju
st fin
ished rifting)SBSC
(10 Ma)
IBM (15
Ma)IBM (10 Ma)
Left-lateral transpression
N
0 200 km~ 8 cm/a
Figure 10. Southwest Japan tectonics from 15 to 10 Ma, from a reconstruction using poles of
rotation of the Philippine Sea plate relative to Eurasia from Gaina and Müller (2007). The gray shaded features are the positions of the Izu-Bonin-Mariana arc (IBM), and Kyushu-Palau ridge (KPR) at 15 Ma, and the solid black line is the position of the Shikoku Basin spreading center (SBSC) at 15 Ma. The dashed-outlined portions of the 15 Ma KPR, SBSC, and IBM approximately show the portions of those features that have been subducted since 15 Ma. The light-gray dashed lines outline the 10 Ma positions of the IBM, SBSC, and KPR. The black arrows (labeled in cm/yr) show approximate Philippine Sea plate–Amurian plate relative rates of motion from 15 to 10 Ma. The present-day east coast Kyushu and Shikoku Island coast line is shown in light gray, while the 15 Ma position of the coast line is in black.to Shikoku Island and Kyushu. Note that the
Shikoku Basin (30–15 Ma; Okino et al., 1999; Sdrolias et al., 2004) and the Sea of Japan (ages for the Sea of Japan opening vary from 30 to 12 Ma: Otofuji and Matsuda, 1983; Tamaki et al., 1992; and Lee et al., 1999; to 16–14 Ma: Otofuji et al., 1991) had largely fi nished open-
ing by 15 Ma, so the development of these basins does not impact our reconstructions.
Tectonic Reconstruction for 15–10 Ma
Between 15 and 10 Ma, the kinematics of the
Philippine Sea plate relative to Eurasia in the southwest Japan region were markedly different from today (according to paleomagnetic studies of Hall et al., 1995a). The portion of the south-west Japan plate boundary adjacent to the Phil-
ippine Sea plate was dominated by left- lateral transpression (Fig. 10). The past kinematics of Philippine Sea plate motion from Hall et al. (1995a) and Sdrolias et al. (2004) requires rapid northeastward migration of the triple junction between southwest Japan, the Philippine Sea plate, and the Pacifi c plates from 15 to 5 Ma.
At 15 Ma, the triple junction was located near Shikoku Island, and by 10 Ma it had migrated northeast ~300 km to a location somewhere between the Kii Peninsula and Boso Peninsula region. East of this triple junction, subduction of the Pacifi c plate is occurring. Thus, if our recon-
struction is correct, prior to 15 Ma, subduction of the Pacifi c plate likely occurred beneath most
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Mahony et al.
2214 Geological Society of America Bulletin, November/December 2011of southwest Japan. At 15 Ma, the Shikoku
Basin spreading center and the Kyushu-Palau ridge were intersecting the subduction margin somewhere offshore the Ryukyu Islands, south-west of Japan.
Sdrolias et al. (2004) suggested that most
of the 25–5 Ma clockwise rotation of the Philippine Sea plate (34° from paleomagnetic studies) occurred between 15 and 5 Ma. If this interpretation is correct, the along-strike migration of the Philippine Sea plate–Pacifi c
plate–southwest Japan triple junction would have been even more rapid than shown here. For example, if 30° of clockwise rotation of the Philippine Sea plate occurred between 15 and 5 Ma (at a rate of ~3°/Myr), the Philip-pine Sea plate–Pacifi c plate–southwest Japan
triple junction would be located well south of Kyushu at ca. 10 Ma. Thus, depending on the details of past Philippine Sea plate rotational kinematics (e.g., Sdrolias et al., 2004), subduc-tion of the Pacifi c plate (and start of Philippine
Sea plate subduction) beneath Kyushu may have initiated as recently as 10 Ma.
Tectonic Reconstruction for 10–5 Ma
Tectonic reconstruction here is for the period
10–5 Ma rather than the similar 10–6 Ma period studied for the volcanism, simply as the fi nite
rotation poles used in this reconstruction (Gaina and Müller, 2007) are listed for 5 Ma, not 6 Ma. The timing of changes in Philippine Sea plate kinematics are not well constrained, so for the purposes of this study the tectonics discussed in this 10–5 Ma period are considered to be com-parable to the 10–6 Ma volcanism.
The large left-lateral component of rela-
tive plate motion at the southwest Japan plate boundary between 15 and 5 Ma continued to cause rapid northeast migration of the posi-tion of the Philippine Sea plate–Pacifi c plate–
southwest Japan triple junction, the Kyushu-Palau ridge, Shikoku Basin spreading center, and the Izu-Bonin-Mariana arc relative to the southwest Japan plate boundary (Fig. 11). By 10 Ma, the Kyushu-Palau ridge was inter-secting the plate boundary offshore southern Kyushu, while the recently extinct Shikoku Basin spreading center intersection point was positioned adjacent to southwestern Shikoku. According to our reconstruction, the Izu-Bonin-Mariana arc point of intersection with the margin at 10 Ma was just east of the Kii Peninsula, and by 8–6 Ma it had migrated close to its current position, adjacent to the Boso Peninsula. However, such a result is at odds with most published literature, which indicates a ca. 15 Ma age for the initiation of Izu-Bonin-Mariana arc collision in its current location in central Japan (Seno and Maruyama, 1984;
128°E 130° 132° 134° 136° 138° 140° 144°22°24°26°28°30°32°34°
142°36°N
~ 7-8 cm/a
~ 8 cm/aoblique subductionoblique subduction
oblique subduction
left-lateral transpression
Philippine Sea PlatePacific
PlateEurasian/Amurian PlateEurasian/Amurian Plate Eurasian/Amurian Plate
KPR (5 Ma
)
KPR (10 Ma)SBSC (5
Ma)SBSC (10 M
a)IBM (5 Ma)
IBM (10 Ma)~ 7 cm/a
N
0 200 kmIzu Bo
nin MarianaArc (10 Ma)
Figure 11. Southwest Japan tectonics from 10 to 5 Ma, based on a reconstruction using poles
of rotation for the Philippine Sea plate relative to Eurasia for this time period from Gaina and Müller (2007). The gray shaded features are the position of the Izu-Bonin-Mariana arc (IBM), Kyushu-Palau ridge (KPR) at 10 Ma, and the solid black line is the position of the Shikoku Basin spreading center (SBSC) at 10 Ma. The dashed-outlined portions of the 10 Ma KPR, SBSC, and IBM approximately show the portion of those features that has been subducted since 10 Ma. The light-gray dashed lines outline the 5 Ma positions of the IBM, SBSC, and KPR. The black arrows (labeled in cm/yr) show approximate Philippine Sea plate–Amurian plate relative rates of motion from 10 to 5 Ma. The present-day east coast Kyushu and Shikoku Island coast line is shown in light gray, while the 10 Ma position of the coast line is in black.Watanabe, 2005; among others; see further dis-
cussion of this issue later in the paper).
Tectonic Reconstruction for 5–2 Ma
Around 5 Ma, Philippine Sea plate motion
changed and began converging in a north-westerly direction relative to southwest Japan (Seno and Maruyama, 1984; Seno, 1989; Hall et al., 1995a). Compared to the 15–5 Ma period, from 5 to 2 Ma there is minimal along-strike migration of the triple junction and the impingement points of the Kyushu-Palau ridge, Izu-Bonin-Mariana arc, and Shikoku Basin spreading center (Fig. 12). Notably, the subduction point of the Kyushu-Palau ridge occurred beneath northern Kyushu throughout
this period, and the Kyushu-Palau ridge col-lision with the subduction margin may have been the cause of anticlockwise rotation of the Kyushu forearc (Wallace et al., 2009a), which started ca. 2–5 Ma (Kodama et al., 1995; Kodama and Nakayama, 1993). Backarc rift-ing in the Okinawa Trough probably began ca. 5 Ma (Sibuet et al., 1987; Wu et al., 2007), so in our reconstruction, the Okinawa Trough has not yet opened prior to 5 Ma. At 5 Ma, at the beginning of rapid anticlockwise rotation of Kyushu, the coastline of Kyushu was also co-linear with the rest of southwest Japan accord-ing to this reconstruction.
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2215Tectonic Reconstruction for 2 Ma–Present
At ca. 2 Ma, the Philippine Sea plate shifted
its subduction direction from northwest to west-northwest (Nakamura et al., 1984). The increase in the westward component of subduc-tion caused the onset of dextral motion on the Median Tectonic Line (Kamata and Kodama, 1994; Itoh et al., 1998; Kamata, 1998). This change in subduction direction also caused the subduction point of the Kyushu-Palau ridge to start migrating southwards (opposite to its previous northwards migration direction), from its location offshore northern Kyushu at ca. 2 Ma, to its present position ~80 km to the south (Fig. 13).Comparison with Previous Reconstructions
of Southwest Japan Tectonics
Our tectonic reconstructions suggest that, prior
to 15 Ma, subduction of the Pacifi c plate occurred
beneath most of southwest Japan and Kyushu. Starting ca. 15 Ma (depending on the details of Philippine Sea plate kinematics), the Kyushu region was straddling the Philippine Sea plate–southwest Japan boundary zone, which was dominated by left-lateral transpression (Fig. 10). From 15 Ma until ca. 5–6 Ma, buoyant Shikoku Basin crust was subducting beneath Kyushu (Figs. 10 and 11); at ca. 5–6 Ma, the older, more deeply subducted west Philippine Basin slab (to
128°E 130° 132° 134° 136° 138° 140° 142° 144°22°24°26°28°30°32°34°36°N
~6 cm/a
(PSP rel
Eurasia) ~5 cm/a
SBSC (5 Ma)
KPR (2 Ma)IBM (2 Ma)SBSC (2 Ma
)
Philippine Sea platePacific
plateEurasian/Amurian plate
N
0 200 kmKyushu
Palau Ridge (5 Ma
)Kyushu Palau Ridge (5 Ma)Kyushu
Palau Ridge (5 Ma
)Izu Bon
in Mar
ianaIzu Bonin M
arianaAr
c (
5 Ma)Arc (5 M
a)Izu Bonin MarianaAr
c (5 Ma)
Figure 12. Southwest Japan tectonics from 2 to 5 Ma, based on a reconstruction using poles
of rotation of the Philippine Sea plate relative to Eurasia from Seno et al. (1993). The gray shaded features are the position of the Izu-Bonin-Mariana arc (IBM) and Kyushu-Palau ridge (KPR) at 5 Ma, and the solid black line is the position of the Shikoku Basin spread-ing center (SBSC) at 5 Ma. The dashed-outlined portions of the 5 Ma KPR, SBSC, and IBM show the portion of those features that has been subducted since 5 Ma. The light-gray dashed lines outline the 2 Ma positions of the IBM, SBSC, and KPR. The black arrows (labeled in cm/yr) show approximate Philippine Sea plate–Amurian plate relative rates of motion from 5 to 2 Ma. The present-day east coast Kyushu and Shikoku Island coast line is shown in light gray, while the 5 Ma position of the coast line is in black.the west of the Kyushu-Palau ridge) began sub-
ducting beneath Kyushu. Around 5–6 Ma, the kinematics of the Philippine Sea plate changed signifi cantly (Hall et al., 1995a), causing the Phil-
ippine Sea plate–southwest Japan plate boundary in the Kyushu region to change from strike-slip dominated to nearly pure convergence, coincid-ing with evidence for reinitiation of subduction-related volcanism in Kyushu ca. 5–6 Ma (Seno, 1989; Kamata and Kodama, 1999; Fig. 12).
Previously published tectonic reconstruc-
tions of southwest Japan give widely varying results. Variation in reconstructions prior to 1995 can be reasonably explained by the lack of paleomagnetic data, which have become avail-able for later studies (e.g., Hall et al., 1995a). Most reconstructions of the southwest Japan region assume that the Izu-Bonin-Mariana arc, Kyushu-Palau ridge, and Shikoku Basin spread-ing center have been generally at the same loca-tion as they are today for the past 15 Myr (e.g., Seno and Maruyama, 1984; Hibbard and Karig, 1990; Jolivet et al., 1994; Taira, 2001; Kimura et al., 2005), or that these features were located even farther to the northeast ca. 15 Ma than they are today (Otsuki, 1990). However, these recon-structions are not consistent with the kinematics of the Philippine Sea plate, as constrained by paleomagnetic and seafl oor spreading studies
(Hall et al., 1995a, 1995b; Hall, 2002; Sdrolias et al., 2004) and seismic tomography studies (Miller et al., 2006). As shown in Figures 10–13, if the estimates of past Philippine Sea plate motion used to constrain the reconstructions shown here are correct, they require the confi g-
uration of the southwest Japan plate boundary to have evolved dramatically since 15–20 Ma, including rapid northeastward migration of the Philippine Sea plate–Pacifi c plate–southwest
Japan triple junction.
Our reconstructions of the southwest Japan
region are consistent with the more regional southeast Asia reconstructions presented by Hall (2002), Sdrolias et al. (2004), and Gaina and Müller (2007), which is expected given that their pre–2 Ma estimates of Philippine Sea plate–Eurasian plate kinematics were used to derive these reconstructions. For example, at 15 Ma, Sdrolias et al. (2004) position the Shi-koku Basin spreading center point of intersec-tion just south of Kyushu, similar to the 15 Ma scenario presented here. An earlier reconstruc-tion by Lee et al. (1999) is among the only studies focused on the southwest Japan region that accounted for the evolving motion of the Philippine Sea plate in a similar way; they also suggested that subduction of the Pacifi c plate
beneath all of southwest Japan prior to 15 Ma and that, for some period after 15 Ma, the southwest Japan–Philippine Sea plate bound-
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Mahony et al.
2216 Geological Society of America Bulletin, November/December 2011ary was strike-slip dominated. Our study
supports Lee et al.’s (1999) reconstruction, which places the 15 Ma Izu-Bonin-Mariana arc impingement point somewhere between Kyushu and Shikoku Island; they suggest this reconstruction is also consistent with evidence for widespread shortening between Kyushu and Korea at this time. Our study also supports Miller et al.’s (2006) reconstruction, which suggested that the Pacifi c plate– Philippine
Sea plate–southwest Japan triple junction was located near the Kii Peninsula at ca. 10 Ma, based on evidence from the subducted Pacifi c
plate slab morphology from seismic tomo-graphic images.IMPLICATIONS OF THESE
RECONSTRUCTIONS FOR VOLCANO-TECTONIC EVOLUTION OF SOUTHWEST JAPAN
Volcanic-Tectonic Evolution from 15 to 10 Ma
Igneous activity is represented by I- and
S-type plutons trending northeast-southwest through central Kyushu (Fig. 14A) and small volcanic remnants. V olcanic centers in south-west Honshu, Shikoku, and on Kyushu erupted the Setouchi high-magnesian andesites from 12 to 15 Ma (Tatsumi, 1983; Tatsumi et al., 2001). High Sr/Y ratios in the Setouchi high-
128°E 130° 132° 134° 136° 138° 140° 142° 144°22°24°26°28°30°32°34°36°N
~7 cm/a~5 cm/a
SBSC (2 Ma)
SBSC (present)KPR (present)IBM
(present)
Philippine Sea platePacific
plateEurasian/Amurian plate
Izu Bonin MarianaArc (2 Ma)
Kyushu Palau Ridge (2 Ma)
N
0 200 km
Figure 13. Southwest Japan tectonics from 2 Ma to present, based on a reconstruction using
poles of rotation of tectonic blocks and plates (including Philippine Sea plate and the Amu-rian plate) from Wallace et al. (2009a). The gray shaded features are the positions of the Izu-Bonin-Mariana arc (IBM), and Kyushu-Palau ridge (KPR) at 2 Ma, and the solid black line is the position of the Shikoku Basin spreading center (SBSC) at 2 Ma. The dashed-outlined portions of the 2 Ma KPR, SBSC, and IBM show the portion of those features that has been subducted since 2 Ma. The light-gray dashed lines outline the present-day position of the IBM, SBSC, and KPR. The black arrows (labeled in mm/yr) show Philippine Sea plate–Amurian plate relative rates of motion from 0 to 2 Ma. The present-day east coast Kyushu and Shikoku Island coast line is shown in light gray, while the 2 Ma position of the coast line is in black.magnesian andesites are interpreted to indicate
subduction of the young hot Shikoku Basin lith-osphere (Kano et al., 1991; Yamaji and Yoshida, 1998; Tatsumi and Hanyu, 2003). Shimoda et al. (1998) argued on the basis of their isotopic data that the Setouchi high-magnesian andesites were generated by interaction of MORB mantle with partial melt of subducted Shikoku Basin sediment at an anomalously high temperature. Both these scenarios are reasonably consis-tent with our tectonic reconstruction (Fig. 10). By 14 Ma, young Shikoku Basin lithosphere was being subducted beneath Kyushu and Shi-koku Island, and also beneath the Kii Penin-sula region by 13 Ma (Fig. 14A). Moreover, if the already subducted Izu-Bonin-Mariana arc had a more northeasterly trend than currently observed (assuming that the subducted Izu-Bonin-Mariana arc had the same orientation as
Figure 14. Volcanic and tectonic evolution
of southwest Japan from 15 Ma to present. Volcanic centers are denoted by black trian-gles with black fi ll (polygenetic volcanism),
no fi ll (monogenetic volcanism), and gray
fi ll (areas of volcanism, e.g., lava-dominated
eruptions). Plutons are marked by solid black circles with either a white I (I-type pluton) or S (S-type pluton), with pluton locations from Kimura et al. (2005). Inset enlarged diagrams of Kyushu volcanic cen-ters in the bottom right of each fi gure rep-
resent the dotted region on the main map. Figures 14A (15–10 Ma), 14B (10–5 Ma), 14C (5–2 Ma), and 14D (2–0 Ma) show southwest Japan tectonics from 15 to 0 Ma, from a reconstruction using poles of rota-tion of the Philippine Sea plate relative to Eurasia from Gaina and Müller (2007). The gray shaded features are the positions of the Izu-Bonin-Mariana arc (IBM) and Kyushu-Palau ridge (KPR) at the oldest end of the age range for each fi gure, and the solid black
line is the position of the Shikoku Basin spreading center (SBSC) at the oldest end of the age range. The dashed-outlined portions of the KPR, SBSC, and IBM approximately show the portions of those features that have been subducted. The light-gray dashed lines outline the position of the IBM, SBSC, and KPR at the youngest end of the age range. The black arrows (labeled in cm/yr) show approximate Philippine Sea plate–Amurian plate relative rates of motion for each age range. The present-day east coast Kyushu and Shikoku Island coast line is shown in light gray, while the position of the coast line at the oldest end of the age range is in black.
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2217128°E 130° 132° 134° 136° 138° 140° 142° 144°22°24°26°28°30°32°34°36°N
~7 cm/aSBSC (2 Ma)SBSC (present)
KPR (present)IBM(present)
Philippine Sea platePacific
plateEurasian/Amurian plate
IBM (2 M
a)
N0 200 km
128°E 130° 132° 134° 136° 138° 140° 142° 144°22°24°26°28°30°32°34°36°N
~6 cm/a
(PSP rel Eurasia) ~5 cm/a
SBSC (5 Ma)IBM (5 Ma)
IBM (2 Ma)
SBSC (2 Ma)
Philippine Sea platePacific
plateEurasia/Amurian plate
N0 200 kmKPR (2 Ma)KPR (5 Ma)128°E 130° 132° 134° 136° 138° 140° 144°22°24°26°28°30°32°34°36°N
142°Left-lateral transpressionPacific
plate
IBM (10 Ma)
N0 200 km
128°E 130° 132° 134° 136° 138° 140° 142° 144°22°24°26°28°30°32°34°36°N
~8 cm/a
~8 cm/aSubductionSubduction
Pacific
plate
Philippine
SeaplateEurasian/Amurian plate
KPR (15 Ma)SBSC (15 Ma)
Left-lateral transpression
N0 200 km
Volcanics key:
Polygenetic volcano
Monogenetic volcanoAreas of lava-dominated eruptionsPlutonIBM (10 Ma)
KPR (10 Ma)
~5 cm/
a ~7-8 cm/a
~8 cm/aOblique subduction
Philippine Sea plateEurasian/Amurian plate
KPR (5 Ma)
KP
R (10 Ma)SBSC (5 Ma)
SBSC (10 Ma)IBM (5 Ma)
~7 cm/a IBM (1
0 Ma)SBS
C
(10 Ma)IBM (1
5 Ma
)
KPR (2 Ma)B A
C DIIIIII
XXI I
S
SS
S
15–10 Ma
5–2 Ma 2–0 Ma10–5 Ma
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Mahony et al.
2218 Geological Society of America Bulletin, November/December 2011observed today), the Shikoku Basin lithosphere
could have begun being subducted beneath the entire southwest Japan region even earlier. Our reconstructed age of onset of Shikoku Basin–lithosphere subduction beneath southwest Japan (ca. 14 Ma) is similar to the age of the erupted Setouchi high-magnesian andesites (13.7 Ma), suggesting the subduction could be a cause of the volcanism. The OIB-type volcanism, which occurred at 15 Ma in the northwestern volcanic region in Kyushu (Shinjo et al., 2000; Uto et al., 2004), suggests that asthenospheric injection occurred beneath this region.
Our reconstructions suggest that the intersec-
tion between the Izu-Bonin-Mariana arc and the southwest Japan plate boundary was located somewhere between Kyushu Island and Shi-koku Island at ca. 15 Ma. Paleomagnetic studies document 45º of clockwise rotation of southwest Honshu in the early to middle Miocene (Otofuji and Matsuda, 1983; Celaya and McCabe, 1987; Otofuji et al., 1991). Note that this time win-dow for the rotation of southwest Honshu is just prior to the time frame (15 Ma) of our recon-struction, and so for simplicity it is not included in our reconstructions. Lee et al. (1999) pro-posed that the Izu-Bonin-Mariana arc collision with Kyushu acted as a pivot point for the rapid clockwise rotation of southwest Honshu during the Sea of Japan opening, consistent with our reconstruction. Rapid arc rotation about a pivot point at a subduction/collision transition is a common process worldwide (e.g., Wallace et al., 2005, 2009b), and the along-strike change from indentation of the Izu-Bonin-Mariana arc at the southwest Japan plate margin (near Kyushu) to subduction of the Pacifi c plate farther northeast
is a plausible mechanism to trigger rapid Mio-cene clockwise rotation of southwest Honshu.
Volcanic-Tectonic Evolution from 10 to 6 Ma
Many workers have noted an absence of
subduction-related volcanism in Kyushu between 10 and 6 Ma (Fig. 14B), interpreting this as a cessation in subduction of the Philip-pine Sea plate offshore Kyushu (e.g., Uto, 1989; Kamata, 1992). Confl icting views are held by
other authors who either suggest there was no late Miocene halt in subduction of the Philip-pine Sea plate (e.g., Maruyama et al., 1997; Kimura et al., 2003), who suggest there was subduction since 7 Ma (Niitsuma, 1988), or who propose an alternative time period for a “stag-nant” phase in subduction (e.g., Taira [2001] suggests 14–8 Ma). Integrated Ocean Drilling Program data support low levels of volcanism in this period (Cambray and Cadet, 1994) with a marked increase starting between 6 and 7 Ma. The proposed halt in subduction is summarized by Kamata and Kodama (1994) and mainly
based on interpretation of the geochemistry of rather sparse volcanic rocks in this period, which lack a slab fl uid signature. Between ca.
15 and 6 Ma, the relative plate motion offshore Kyushu on the Nankai Trough–Ryukyu Trench is dominated by strike slip with some slow con-vergence (Fig. 14B), suggesting that subduction was much slower from 10 to 6 Ma (and perhaps even longer, from ca. 14 to 6 Ma, similar to the suggestion of Taira, 2001) than it is today.
Our reconstructions highlight an additional
explanation for the apparent cessation of subduction-related volcanism from 10 to 6 Ma: the presence of the young, buoyant Shi-koku Basin lithosphere that was probably being subducted obliquely at a shallow angle beneath Kyushu during this time (Fig. 14B). In modern-day southwest Honshu, the Shikoku Basin lithosphere is being subducted at a very shallow angle, with the leading edge of the plate being currently at 70 km depth (Fig. 2), and an absence of active subduction-related volcanism. The modern setting in southwest Honshu thus provides an analog for Kyushu at 10–6 Ma, where the recently formed Shikoku Basin litho-sphere was subducted in a similar fashion.
In spite of this proposed halt in subduction,
low-intensity volcanism continued throughout this period in the western parts of Kyushu at Iki, Kitamatsuura, Shimoshima, and Nansatsu (see Fig. 6 for locations). V olcanism at Iki Island is consistently characterized by OIB-type geo-chemistry through time, refl ecting its backarc
setting (Shinjo et al., 2000). The episodic Kita-matsuura basalt volcanism has OIB and alkaline characteristics, but the youngest unit at 6 Ma is calc-alkaline (Sakuyama et al., 2009). The physi-cal shape of Kyushu at 6 Ma was very different to now. Notably, prior to the opening of the Oki-nawa Trough, the Kitamatsuura and Nansatsu regions would have been geographically closer (e.g., Fig. 11 shows the reconstructed geography of Kyushu at 6 Ma). The calc-alkaline nature of the 6 Ma Kitamatsuura basalts and the island arc basalt volcanism in the Nansatsu area of south-ern Kyushu from 6.4 to 5.9 Ma (Fig. 6), together with the ODP observations, indicate onset of subduction-related volcanism at ca. 6.5 Ma. This suggests that the older, more steeply dipping West Philippine Basin portion of the Philippine Sea plate had migrated to the Kyushu region (consistent with our reconstructions; Fig. 14B) and had been subducted deeply enough to initiate the generation of island arc–type magmas in the Nansatsu and Kitamatsuura regions. Sakuyama et al. (2009) suggested that the change from OIB-type alkali to calc-alkaline basalts in the Kitamatsuura region was due to the shallow-ing of the melt-extraction depth in an upwelling mantle plume on the basis of their petrological
data. However, when the volcanic-tectonic sys-tem is reconstructed as a whole, we suggest that the Nansatsu and Kitamatsuura volcanic areas were part of the same volcanic system. The calc-alkaline Kitamatsuura volcanism was followed a few million years later by the Arita rhyolite vol-canoes, representing the next stage in the volca-nic evolution of the region, with focusing of the volcanism into more mature volcanic centers.
Volcanic-Tectonic Evolution from 6 to 2 Ma
Subduction of the Philippine Sea plate in a
north-northwest direction, which began at 6.5 Ma, had a major impact on the volcanic and tectonic events that followed. By 6 Ma, the Kyushu-Palau ridge had migrated to northeast Kyushu, and remained there until 2 Ma (Fig. 14C). We propose that the Kyushu-Palau ridge subduc-tion adjacent to northeast Kyushu was the cause of the 30º anticlockwise rotation of the Kyushu forearc documented from paleomagnetic stud-ies during this time (Kodama et al., 1995). This rotation was initiated by an along-strike change from subduction of the shallowly dipping Shi-koku Basin lithosphere and Kyushu-Palau ridge, to subduction of the older denser West Philippine basin lithosphere, which was subducted steeply and likely rolled back farther south. Together these competing forces created a torque that gen-erated rotation of the Kyushu forearc. Rollback of the West Philippine Basin slab adjacent to the Ryukyu Trench and Kyushu, and rotation of the Kyushu forearc, has infl uenced the kinematics
of north-south–directed extension in the Beppu-Shimabara graben (including the Hohi volcanic zone) of central Kyushu.
Coincident with the onset of extension, the
Hohi volcanic zone was the site of major erup-tive outpourings from 6 to 2 Ma, with a peak in production at 5 Ma. During this time period, the Kyushu-Palau ridge subduction point under-went minimal along-strike migration due to the nearly orthogonal convergence between the Philippine Sea plate and Eurasia (Fig. 14C). We suggest that the voluminous volcanism in the Hohi volcanic zone during this period is related to a combination of upper-plate exten-sion and Kyushu-Palau ridge subduction. The subduction of the Kyushu-Palau ridge directly beneath northeast Kyushu at 5 Ma introduced a large amount of fl uids into the melt hot zone
feeding the Hohi volcanic zone, which trig-gered arc magmatism. V olcanic responses to the reinitiation of subduction are confi rmed by an
observed increase in K content in Hohi volcanic zone magmas from 6 Ma onwards (Nakada and Kamata, 1991), confi rming an increased fl uid
signature and subducted slab infl uence.
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Geological Society of America Bulletin, November/December 2011 2219V oluminous lava plateaus were produced in
areas corresponding to present-day Aso and Aira calderas. One characteristic of Kyushu vol-canism appears to be voluminous andesite lavas in an area prior to caldera-forming eruptions involving more silicic magmas (part of the cal-dera cycle). It is suggested that the magma intru-sion into the lower crust associated with these precursory lavas serve to thermally weaken an area (Annen et al., 2006), and lead to conditions conducive to caldera formation.
During the 6–2 Ma period, the northwestern
region volcanoes erupted OIB-type basalt, unin-fl uenced by subduction components. The initia-
tion of subduction and arc processes perhaps led to the pulse of backarc volcanism; OIB-type volcanism occurred around Fukuoka from 4 to 1 Ma (Hoang and Uto, 2003; Fig. 8).
Volcanic-Tectonic Evolution from 2 Ma
to Present
The subduction direction of the Philippine Sea
plate beneath southwest Japan at 2 Ma changed from north-northwest to northwest (Yamaji, 2003; Fig. 14D), causing signifi cant volcanic
and tectonic consequences such as initiation of dextral strike slip on the Median Tectonic Line (Itoh et al., 1998). This change in plate motion also caused the migration direction of the Kyushu-Palau ridge subduction point along the Nankai Trough–Ryukyu Trench to change from northeast to southwest (the Kyushu–Palau ridge subduction point currently migrates southwest at 40 km/Myr; Wallace et al., 2009b). We contend that southwestward migration of the buoyant, shallowly subducting Shikoku Basin lithosphere into the northern Kyushu region since 2 Ma has been the cause of lithospheric shortening in the Hohi volcanic zone since 0.7 Ma and cessa-tion of Hohi volcanic zone graben development (Kamata et al., 1988).
The rate of volcanism in the Shimabara gra-
ben (near Unzen volcano) has increased since 6 Ma (Yokose et al., 1999), supporting the idea that increasing extension in the region related to rollback of the Philippine Sea slab offshore Kyushu (Okinawa Trough) has facilitated a greater rate of volcanism. Rollback of the Philip-pine Sea plate has also impacted the location of volcanism in southern Kyushu (Fig. 14D). From ca. 6 to 1 Ma, the volcanic front was located near the Nansatsu, Hokusatsu, and Hisatsu volcanic centers; however, increased Philippine Sea plate slab rollback shifted the volcanic front eastward to its current location in the Kagoshima graben (Yamaji, 2003; Figs. 8 and 9).
In the past one million years, a number of
large caldera volcanoes (e.g., Aira, Aso, etc.; see Fig. 9) have formed in Kyushu along the volcanic front. The two major areas of caldera
formation in Kyushu appear to coincide with areas of extensional normal faulting, namely the Kagoshima graben and the Beppu-Shimabara graben (Fig. 3). The fi rst of these calderas to
form was the Shishimuta caldera from 1 to 0.9 Ma (Fig. 9). Shishimuta caldera demon-strates how subtle changes in the tectonics of a region may affect the related volcanism. From 1 to 0.7 Ma, a north-south extensional stress fi eld
was present in the Hohi volcanic zone (reported by Kamata et al., 1988), during which time the large Shishimuta caldera formed (Kamata et al., 1988). At 0.7 Ma, a local, weak compressional stress regime initiated in the Hohi volcanic zone (Kamata et al., 1988; Itoh et al., 1998), and since that time only lava dome volcanoes have formed in the Hohi volcanic zone. This indicates a tec-tonic relationship with the style of volcanism, and suggests that other large calderas (i.e., Aso, Aira, Ata, Kakuto, Kobayashi, and Kikai) may have formed in periods of local extension within their respective grabens. The alternative possi-bility in the volcano-tectonic “chicken-and-egg” conundrum is that intense magmatism associ-ated with calderas and underplating weakened the upper plate to induce extension. This sug-gests that grabens preferentially form in areas of high magma fl ux, such as the extensive lava pla-
teaus that occurred in Kyushu prior to calderas.
The location of calderas could be explained
by the positions of physical features (such as the Kyushu-Palau ridge) on the subducting Philippine Sea plate. The subduction of sea-mounts and ridges introduces a large amount of fl uid-rich minerals into the subduction zone
as well as tectonically eroded sediments (Bangs et al., 2006). Once the slab reaches a depth of ~100 km, these fl uid-rich sediments induce the
formation of hydrous island arc–type magmas, leading to explosive volcanism (England et al., 2004). Recent Aso and Kirishima basalts both show signifi cantly higher B/Nb and B/Zr
(Fig. 4) than the northern volcanic front basalts (Oninomi, Yufu, and Kuju). In addition, the hornblende-bearing silicic products from the northern volcanic front show high Sr/Y ratios and adakite composition (Fig. 4). The adakitic magmatism in the northern volcanic front prob-ably corresponds to the subduction of the young (and hot) Shikoku Basin. The source mantle of Aso and Kirishima basalts may be metaso-matized by the modern slab-derived fl uid from
the subducted old West Philippine basin and/or Kyushu-Palau ridge. Higher ratios of B/Nb and B/Zr are noted at Kirishima volcano than at Aso, supporting the idea of the Kyushu-Palau ridge migrating southwestwards along the Nankai Trough–Ryukyu Trench since 5–6 Ma. The combination of the position of the sub-ducted Kyushu-Palau ridge with upper-plate
extension is seen to dramatically increase the volume of magma produced. For example, the backarc extension termination point coincides with the Kyushu-Palau ridge subduction point near Aso volcano. Shishimuta caldera, or even the whole Hohi volcanic zone located slightly northeast of Aso, may represent an older ver-sion of this extensional and fl uid-rich subduc-
tion environment. An example of this scenario is seen in Kamchatka at the Kluchevskoy volcano complex, which is one of the most productive volcanoes in the world. Kluchevskoy volcano complex sits in a 200-km-wide graben structure located above where the Emperor Seamount chain (on the Pacifi c plate) subducts beneath
the Eurasian plate (Dorendorf et al., 2000). The high magma production rate and enriched
18O
isotopes were suggested by Dorendorf et al. (2000) to be a consequence of the combination of high fl uid input over time due to seamount
chain subduction, and intra-arc extension where the seamount chain is subducted. The examples from Kyushu and Kamchatka demonstrate that volcanism is strongly infl uenced by the age and
composition of subducted lithosphere.
Aso caldera has had four major caldera-
forming episodes since 0.3 Ma, demonstrating common cyclic behavior for a large caldera system. Cyclic volcanism at calderas is usually attributed to magma fl uxing, magma differentia-
tion, and episodic generation of silicic magma chambers (e.g., Rytuba, 1994; Riciputi et al., 1995; Troll et al., 2002), so no specifi c tectonic
controls are needed to explain the episodic caldera behavior of Aso volcano. However, an alternative idea for Aso relates to the fact that the Kyushu-Palau ridge appears to subduct directly beneath it (Fig. 1). Bathymetric images of the nonsubducted part of the Kyushu-Palau ridge (e.g., Fig. 1) show that the Kyushu-Palau ridge is not a smooth continuous feature but lumpy with changes in width and height along its length. Subduction of this “lumpy” remnant arc would cause an uneven fl uid fl ux into the
mantle, leading to periods of highly fl uid-rich,
voluminous volcanism, as seen at Aso. We have discussed many instances of volcanic-tectonic evolution and interactions in Kyushu; however, there are still aspects of Kyushu’s volcanic set-ting that are poorly understood, for example, the nonvolcanic southeastern region between Aso volcano and Kirishima volcano.
IMPLICATIONS FOR TIMING OF
IZU ARC COLLISION WITH CENTRAL JAPAN
Most published tectonic histories of Japan
assume a mid-Miocene onset of collision of
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Mahony et al.
2220 Geological Society of America Bulletin, November/December 2011the Izu Peninsula with central Japan (Seno and
Maruyama, 1984; Hibbard and Karig, 1990; Taira, 2001; Kimura et al., 2005). Such a model requires the Izu collision to have been ongo-ing in its current location since ca. 15 Ma, and assumes that the confi guration of tectonics in
southwest and central Japan has changed little since that time. The primary evidence support-ing a mid-Miocene timing for the Izu collision is the intrusion of the 15.7–7.4 Ma Kofu Granitic Complex (which is thought to be derived from the Izu arc) into the Cretaceous–Paleogene Shi-manto belt (Kawano and Ueda, 1966; Shibata et al., 1984; Saito and Kato, 1996; Saito et al., 1997). The Shimanto Belt is an ancient accre-tionary complex and one of the primary bedrock terranes in central Japan in the region of the Izu collision. The other piece of evidence cited in support of a middle Miocene Izu collision is the mid-Miocene age of trough-fi ll sediments adja-
cent to the Koma-Kushigatayama block of the Izu collision zone (Aoike, 1999). The assump-tion that the Izu collision began in the middle Miocene has dominated the majority of pub-lished Miocene to present tectonic interpreta-tions of Japan.
However, if the past kinematics of the Philip-
pine Sea plate, as determined from paleomag-netic data and seafl oor spreading studies (Hall
et al., 1995a, 1995b; Hall, 2002; Sdrolias et al., 2004), are correct, the Izu arc collision with central Japan should have occurred much more recently, probably ca. 6–8 Ma, as suggested by the reconstructions of Lee et al. (1999) and Toda et al. (2008). Our age for onset of the Izu collision is consistent with the inferred age of the Tanzawa block (part of the Izu arc) col-lision at 6.8 Ma (Yamamoto and Kawakami, 2005). These reconstructions differ markedly from most published reconstructions of Japa-nese tectonics (e.g., Seno and Maruyama, 1984; Hibbard and Karig, 1990; Jolivet et al., 1994; Taira, 2001; Kimura et al., 2005; Yamazaki et al., 2010, among others), many of which were developed prior to the collection of paleomag-netic and other data sets documenting the past history of the Philippine Sea plate. As discussed in the previous section, our reconstructions explain many features of the volcano-tectonic evolution of the Kyushu region that are diffi cult
to understand in the context of most published reconstructions of southwest and central Japan.
We suggest that the tectonic and volcanic his-
tory of Kyushu and southwest Japan supports the plate reconstructions presented here, for sev-eral different reasons.
(1) Prior to now, there has been no explanation
for the hiatus in subduction-related volcanism in Kyushu between 10 and 6 Ma (Cambray and Cadet, 1994; Kamata and Kodama, 1999). Our reconstructions suggest that highly oblique sub-
duction of the newly formed, shallowly subduct-ing Shikoku Basin provides a likely explanation for this. Under this model, Kyushu from 10 to 6 Ma (Fig. 11) is analogous to modern-day south-west Honshu where the buoyant Shikoku Basin is currently being subducted at a low angle and the leading edge of the Philippine Sea plate is at ~70 km depth (Fig. 2), and there is currently no active arc volcanism (Fig. 14B). If the Shikoku Basin was being subducted beneath Kyushu from 10 to 6 Ma, as our reconstruction and the lack of subduction related volcanism suggests, the Izu arc would have been intersecting the plate bound-ary ~100–200 km southwest of its current loca-tion for much of this time period.
(2) Middle Miocene (16–14 Ma; Otofuji et al.,
1991) estimates for the most rapid opening of the Sea of Japan are incompatible with the onset of the Izu collision ca. 15 Ma. If the Izu collision began in its current location in central Japan at 15 Ma, the newly formed Shikoku Basin crust and the buoyant Izu-Bonin arc would have been subducted beneath southwest Japan adjacent to the area of most rapid Sea of Japan opening dur-ing rifting. For backrifting to occur in the Sea of Japan, the slab being subducted beneath south-west Japan must have been able to roll back. It is highly unlikely that the Shikoku Basin crust and the Izu arc were converging on southwest and central Japan at the time of the opening of the Sea of Japan, because these buoyant features would have retarded convergence at the trench and resisted slab rollback and rapid backarc opening in the Sea of Japan. For mid-Miocene rifting in the Sea of Japan to occur, it is more likely that the older, denser Pacifi c plate was sub-
ducting beneath most of southwest Japan at that time (and that the Philippine Sea plate–southwest Japan–Pacifi c plate triple junction was located
near Kyushu at ca. 15 Ma), consistent with our reconstructions shown in this paper (Fig. 10).
(3) Approximately 45° of clockwise tec-
tonic rotation of southwest Honshu occurred in the middle Miocene, simultaneous with the Sea of Japan opening. The kinematics of rota-tion of southwest Honshu and the simultane-ous opening of the Sea of Japan are consistent with rotation of southwest Honshu about a pivot point located near Kyushu. Collision of the Izu-Bonin-Mariana arc with Kyushu is expected just prior to 15 Ma in our reconstructions, and we suggest that the collision formed a pivot point for the clockwise rotation of southwest Honshu. All modern-day examples of rapid tectonic rota-tion of arcs are associated with an along-strike transition from collision to subduction (Wal-lace et al., 2005, 2009b). In many cases these rotating arcs are also associated with backarc rifting, and collision of a buoyant feature with the subduction margin acts as a pivot point for
the rotation of the arc (Wallace et al., 2005). The kinematics of rotation of southwest Honshu and the Sea of Japan opening are very similar to these modern-day examples, and it is likely that a similar scenario has occurred in southwest Japan (e.g., rapid arc rotation about a collisional pivot point). Lee et al. (1999) suggested that a wide zone of early to middle Miocene shorten-ing between the Korean Peninsula and south-west Japan, near Tsushima Island, provides evi-dence for a collisional pivot point during the Sea of Japan opening.
The history of migration of the southwest
Japan–Philippine Sea plate–Pacifi c triple junc-
tion along the plate boundary is central to under-standing the evolution of Japanese tectonics and volcanism since the Miocene. The fundamental disagreement between the timing of Izu arc collision with central Japan and the history of motion of the Philippine Sea plate clearly needs to be resolved, as this has major implications for our understanding of the history of triple junction migration along the plate boundary. It is obvious that the data constraining the history of Philippine Sea plate motion, the evidence for timing of Izu arc collision, and timing of back-arc rifting in the Sea of Japan require serious reevaluation. However, we contend that recon-structions accounting for current knowledge of the past motion of the Philippine Sea plate (e.g., Hall et al., 1995a, 1995b; Sdrolias et al., 2004) are more consistent with the tectonic and volca-nic evolution of Kyushu and southwest Japan.
CONCLUSIONS
The diverse volcanism that has occurred in
Kyushu since the early Miocene can largely be explained as a consequence of evolution of the plate-boundary confi guration in the south-
west Japan region. In the plate tectonic recon-structions shown, the subduction of the Pacifi c
plate occurred beneath most of southwest Japan prior to 15 Ma, and volcanism before this time is likely related to subduction of the Pacifi c
plate. From 15 to 5 Ma, relative plate motion at the Philippine Sea plate–southwest Japan boundary is highly oblique (dominated by left-lateral strike slip), and we suggest that a lack of subduction-related volcanism from 10 to 6 Ma can be explained by (1) a smaller convergent component of relative plate motion and (2) low-angle subduction of young, buoyant Shikoku Basin lithosphere, similar to what is occurring in southwest Honshu today.
A change in Philippine Sea plate motions ca.
5 Ma led to more rapid, nearly trench-normal convergence, and subduction of the Eocene– Oligocene west Philippine Basin beneath
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Volcano-tectonic interactions during rapid plate-boundary evolution in the Kyushu region, SW Japan
Geological Society of America Bulletin, November/December 2011 2221Kyushu. This helps explain the increasingly
arc-like signature of volcanism in Kyushu since 6 Ma. The voluminous recent volcanism in the Aso and Hohi volcanic zone regions can be explained by a combination of the local exten-sional tectonics and subduction of the fl uid-rich
Kyushu-Palau ridge beneath both the Hohi vol-canic zone from 2 to 6 Ma and the Aso volcano area since 2 Ma. A strong spatial correlation between productive volcanism and extensional tectonics in Kyushu suggests that the occur-rence of tectonic extension can help to enhance the rate of volcanism and vice versa. Geochemi-cal characteristics and voluminous volcanism in the Hohi volcanic zone and near Aso volcano can be explained by upper-plate extension in these regions and subduction of the fl uid-rich
Kyushu-Palau ridge.
ACKNOWLEDGMENTS
We express our utmost thanks to Mark Cloos and
Chuck Connor for their valuable input and comments, the Nuclear Waste Management Organization of Japan (NUMO) for funding much of this work, Neil Chap-man for encouraging our ideas, and Nicholas Barnard for his comments. Sparks acknowledges a European Research Council advanced grant.
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Interaction of a spreading ridge and an accretionary prism:
Implications from MORB magmatism in the Hidaka magmaticzone, Hokkaido, Japan
Jin’ichirou Maeda DepartmentofEarthandPlanetarySciences,GraduateSchoolofScience,HokkaidoUniversity,Sapporo,
Hokkaido060,Japan
Hiroo Kagami InstituteforStudyoftheEarth’sInterior,OkayamaUniversity,Misasa,Tottori682-01,Japan
ABSTRACT
In the Hidaka magmatic zone of central Hokkaido, Paleogene mafic plutons have
intrudedaLateCretaceoustoearlyPaleogeneaccretionaryprism,resultinginhigh-tem-peraturemetamorphismandanatexis.Primitivebasalticrocks,whichmakeuppartofthese mafic plutons, are typical normal mid-ocean ridge basalt (MORB). Because thismagmatismisprobablysynchronouswitharrivaloftheKula-Pacificridge,weproposethattheHidakamagmatismrepresentsridge-trenchcollision.Isotopicratiosvarysignificantlywithin the mafic plutons, indicating that intracrustal evolution of normal MORB wasaccompaniedbyassimilationoftheaccretionaryprism.Thus,juvenilecontinentalcrustintheHidakamagmaticzoneformedinanear-trenchsettingthroughtheinteractionofaspreadingridgeandanaccretionaryprism,inmarkedcontrasttonormalmid-oceanridges.
INTRODUCTION
Ridgesubductionshaveplayedanimpor-
tant role in tectonics, metamorphism, andmagmatismattriplejunctions.Inancientac-cretionaryprisms,near-trenchmagmatism,high-temperature metamorphism, and/oranatexisunrelatedtoarcmagmatismhavebeendocumented(e.g.,Hudson,1979;Sis-sonandPavlis,1993),andsothemostfa-vorable model is ridge subduction (e.g.,Marshak and Karig, 1977). However, thecharacteristicsofridgesubductionarenotwellunderstoodbecausetherearefewop-portunitiesfordirectobservationofdeeperprocesses in both ancient and modernexamples.GEOLOGICOUTLINEOFTHE
HIDAKAMAGMATICZONE
The300-km-longHidakamagmaticzone
cropsoutincentralHokkaido,northernJa-pan(Maeda,1990;Fig.1A).Upthrustingofthesouthernhalfofthezonealongthewest-ernfoothillsoftheHidakaMountainssincethe late Miocene (Kimura, 1986) has ex-posed a section of upper mantle to uppercrust that shallows toward the east (Ko-matsuetal.,1994).
Inthewesternmostpart,upper-mantlepe-
ridotites(e.g.,Horomanperidotitemassif)areinfaultcontactwithcrustalrocks.ThesupracrustallithologiesoftheHidakamag-matic zone are terrigenous sedimentaryrocksandgreenstonesoftheHidakaSuper-
group of latest Cretaceous to early Pa-leogene age (Kiminami et al., 1992) andtheir metamorphic equivalents. Metamor-phicgradeincreasestowardthewestfromgreenschist to granulite facies, and maxi-mumpressureandtemperatureconditionswere 0.7 GPa and 830 8C (Osanai et al.,
1991).Inthegranulitezone,graniticrocks(mostlytonalites)resultedfromanatexisofmetamorphicrocksat55Ma(Owadaetal.,1992). Large gabbro plutons are also ex-posedinthegranulite-andamphibolite-fa-cies zones (Maeda, 1990). High-tempera-ture metamorphism and anatexis probablyresultedfromtheintrusionoflargevolumesofgabbro(MaedaandKagami,1994;Ko-matsuetal.,1994),whichconstituteapprox-imately30%ofthetotalexposureofcrys-talline rocks. Thus, gabbro intrusionprobablyoccurredatabout55Ma.
PLUTONSINTHENORTHERN
HIDAKAMOUNTAINS
InthenorthernHidakaMountains,thick
sectionsofmaficplutonicrocksareexposed(Fig. 1B). From the base, these are theMagarisawa peridotite body, Pankenushi
Figure 1. A: Distribution of plu-
tons (shaded areas). Star marksHoroman peridotite massif. QVFisQuaternaryvolcanicfront.Be-foreopeningofKurilbasininMio-cene time, East Hokkaido block(EHB) was located north of itspresent position, and centralHokkaidofacedPacificOceantoeast.Dashedlineswithtrianglesarepaleo–Japanandpaleo–Kuriltrenches at that time. B: Simpli-fied geologic map of northernHidaka Mountains. Patterns:1—Metabasites of Poroshiriophiolite, 2—peridotites (Ma-garisawa and other bodies),3—pelitic and mafic granulitesand anatexites, 4—layered oli-vinegabbros(Pankenushisuite),5—heterogeneous gabbros anddiorites (Memurodake suite),6—granitic rocks, 7—metasedi-mentary and sedimentary rocks(Nakanogawa Group).
Geology;January1996;v.24;no.1;p.31–34;3figures;2tables. 31
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layered gabbroic intrusion, Memurodake
plutoniccomplex,andseveralgranitecom-plexes. Metamorphic and anatectic rocksderivedfromtheHidakaSupergroup(Na-kanogawa Group; Nanayama et al., 1993)are exposed between and around theseplutons.
TheMagarisawaperidotitebodyiscom-
posedmainlyofplagioclaselherzolite,withintercalatedlayersofolivinegabbro.Petro-graphically, the Magarisawa peridotite issimilar to the Main harzburgite-lherzolitesuite(Takahashi,1991)oftheHoromanpe-ridotite massif. The Pankenushi layeredgabbroic intrusion consists of olivine-richtroctolite,olivinegabbro,andferrogabbro,withintercalationsofanorthosite.Alackofa hiatus in olivine crystallization suggeststhatthePankenushigabbrocrystallizedun-derrelativelyhighpressureconditionscom-patiblewiththoseoftheunderlyinggranu-lites. Although troctolite is present in thelower part and ferrogabbro in the upper-most part, cryptic mineral compositionalvariationsthroughthesequenceareconsist-entwithperiodicreplenishmentofprimitive
magma.TheChirorobasalticdikes,whichareafewcentimetrestoseveraltensofcen-timetresthick,intrudethePankenushigab-bros.Theyhavenochilledmargins,suggest-ing that they intruded before completesolidificationofthegabbros.ThemodeofoccurrenceoftheChirorodikesimpliesthattheyarefeederstothePankenushimagmachamber.TheMemurodakeplutonconsistsof hornblende-bearing gabbro and horn-blende-andbiotite-bearingdiorite.Inout-crop,theMemurodakesuiteisaheteroge-neousassemblageoffine-grainedmaficandcoarse-grained felsic rocks. Although the
Memurodakerocksplotinthecalc-alkalicfieldonsomediscriminationdiagrams,theirdecreasingSiO
2andincreasingFeO*(total
FeasFeO)andTiO2withincreasingFeO*/
MgOduringtheearlystageofevolutionre-sembletholeiiticfractionationtrends.
GEOCHEMISTRYOFPRIMITIVE
MAGMAANDFEATURESOFITSMANTLESOURCE
The Chiroro basalts have a nonporphy-
ritic and/or noncumulus, fine-grained tex-ture, indicating that they represent liquid
32 GEOLOGY,January1996
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compositions.TheirhighMgO,Ni,andCr
contents indicate that they represent un-fractionated, mantle-derived magma (Ta-ble 1). Sr and Nd isotope ratios (Table 2and Fig. 2A), chondrite-normalized rareearth element patterns, and MORB-nor-malizedtraceelementpatterns(notshown)oftheChirorobasaltaresimilartonormalMORB.TheChirorobasaltsalsoplotwithinthenormalMORBfieldinsomediscrimi-nation diagrams. Bulk rock compositionsplottedinthepseudoliquidusphasediagramofFalloonandGreen(1988)indicatethatthe Chiroro magma segregated from rela-tivelyfertileperidotiteatabout1.7GPa.SrandNdisotopicratiosoffertileperidotiteintheHoromanmassif(e.g.,YoshikawaandNakamura, 1994) and intercalated maficlayers in the Magarisawa peridotite body(Maeda and Kagami, 1994) suggest thattheseperidotitesarepossiblemantle-sourcecandidatesfortheChirorobasalt.INTRACRUSTALEVOLUTIONOF
NORMALMORBMAGMAANDORIGINOFARCLIKESIGNATURE
ThePankenushiandMemurodakesuites
show a similar range in Sr and Nd initialisotopicratios(Fig.2A),implyingthatbothsuitesshareacommonpetrogenetichistory.Furthermore,closed-systemevolutionisun-abletoexplaintheisotopicdata.AlthoughthemajorandtraceelementsoftheChirorobasaltcannotbedirectlycomparedtothoseof the Pankenushi and Memurodake gab-bros,theisotopicratiosareidenticaltothedepletedendofbothsuites,suggestingthattheChirorobasaltrepresentstheirparentalmagma.
TheisotopicdiversityofthePankenushi
andMemurodakerocks(Fig.2A)fromnor-mal-MORB–liketoarclike(orbetweentheChirorobasaltand[meta]sedimentaryrocksand/oranatexite)canbeexplainedbyAFC(assimilation and fractional crystallization:DePaolo,1981).AlthoughthePankenushi
gabbrosclusteratthelow-Ndend,theval-ues in the Memurodake rocks are widelyscattered(Fig.2B),indicatingthatthecom-positionofassimilatedrocksand/orratioofassimilation to crystallization was not uni-form in this large magma chamber. Thus,thesignificantisotopicdiversityinthesetwosuitesisnotduetomantle-sourceheteroge-neity,buttointeractionbetweenmantle-de-rivednormal-MORBmagmaandtheaccre-tionary prism. Thus we conclude that theHidakamagmaticzonewasnotamagmaticarc.
DISCUSSIONANDCONCLUSIONS
The Nakanogawa Group comprises an
accretionary prism either along the north-trendingpaleo–Japantrench(Kiminamietal., 1992) or at the junction of the pa-leo–Japan arc-trench system and the east-trending paleo–Kuril arc-trench system(Nanayamaetal.,1993).ThePaleoceneNa-kanogawa Group (Nanayama et al., 1993)immediately predated presumed Eocenenormal-MORBmagmatism,somagmatismmustberestrictedtoanear-trenchposition.The timing of normal-MORB magmatismanddepositionoftheNakanogawaGroupalso coincides with the subduction of theKula-PacificridgeinlatePaleocenetoearlyEocene time (Kimura and Tamaki, 1986).ByrneandDiTullio(1992)showedamoresoutherlypositionoftheKula-Pacificridgeduring the early Paleogene. However,northeastwardmigrationoftheKula-PacificridgesubductionalongtheeasternEurasianmargin is compatible with the geology ofsouthwest Japan, such as the eastwardyoungingofbothin-situMORBvolcanismintheaccretionaryprism(Kiminamietal.,1993) and granite magmatism (Kinoshita,1995).Thus,wearguethattheHidakanor-malMORBwasrelatedtotheKula-Pacific
Figure2.A: «Ndvs.«Srat55Ma.Valuesfor «Ndand «Srarecalculated
withrespecttoachondriticuniformreservoir(CHUR)withpresent
143Nd/144Nd50.512638,147Sm/144Nd50.1966,87Sr/86Sr50.7045,
and87Rb/86Sr50.0827; l147Sm 56.54 310212/y,l87Rb51.42 3
10211/y.B:143Nd/144Ndvs.Ndcontent.SymbolsasinA.Alsoshown
are calculated evolution curves of Chiroro basalt. SFC—simplefractional crystallization, AFC—fractional crystallization with as-similationofgranulite(90%pelitic 110%mafic), r—ratioofassim-
ilation to crystallization rates, SMX—simple mixing with peliticgranulite. Fractionating assemblages in SFC and AFC are olivinefrom 0% to 20% solidification, olivine (30%) 1plagioclase (70%)
from 20% to 35% solidification, and olivine (5%) 1plagioclase
(45%) 1clinopyroxene (50%) from 35% to 100% solidification,
based on modal analysis of Pankenushi gabbros. Partition coeffi-cients of Nd for olivine, plagioclase, and clinopyroxene are 0.006,0.090, and 0.230, respectively.
Figure 3. Magmatic and metamorphic pro-
cessesinHidakamagmaticzoneandschemeforgenerationofcontinentalcrustinfore-arcsettings by ridge subduction.
GEOLOGY,January1996 33
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ridge,whichcollidedwiththeaccretionary
prismbeforethesouthwardopeningoftheKurilbasin(Fig.1A).IfthereconstructionofByrneandDiTullio(1992)isapplicable,some segments of the Kula-Pacific ridge,whichextendedtothenorth,mayhaveen-counteredtheEurasianmargin.
Inaschematicmodelformagmaticand
metamorphicprocessesintheHidakamag-maticzone(Fig.3),normal-MORBmagmasegregated from upwelling asthenosphericmantle along the Kula-Pacific ridge, thenmigratedupwardintothebaseoftheaccre-tionaryprismandformedamagmachamberresulting in magmatic underplating in theforearc.Primitivemagmaswererepeatedlyinjected into the magma chamber, trans-porting heat from the asthenosphere intothe accretionary prism. Fore-arc accretedmaterialwasmetamorphosedatuptogran-ulitefaciesandpartiallymeltedtoformgra-niticmagmas.Normal-MORBmagmafrac-tionated and assimilated the accretionaryprismtoformlayeredolivinegabbroofthePankenushi suite. Sporadic injection ofmagmafromthemagmachamberintotheoverlying metasedimentary rocks and/ortheiranatecticmeltsformedtheheteroge-neous,hybridcomplexoftheMemurodakesuite.
MORB-sediment interaction has been
documentedfromridge-trenchcollisionar-eas, such as western California (JohnsonandO’Neil,1984;Sharmaetal.,1991;ColeandBasu,1995)andsouthernChile(Kae-ding et al., 1990). In the Shimanto accre-tionarycomplexofsouthwestJapan,wheretheShikokubasinspreadingridgeandNan-kai trench collided (Hibbard and Karig,1990a),MORB-likegabbros(Miyake,1985;Hibbard and Karig, 1990b) and anatecticgranites(Murata,1984)havebeenreported.In the Chugach metamorphic complex ofAlaska, high-temperature metamorphicrocks and anatexites are exposed in a Pa-leogeneaccretionaryprism(e.g.,SissonandPavlis, 1993). We propose here that thepresenceofMORBmagmatism,high-tem-peraturemetamorphism,andanatexisoftheaccretionaryprismintheHidakamagmaticzoneisconsistentwithridgesubduction.
Magmatic and metamorphic processes
whenaspreadingridgeinteractswithanac-cretionary prism are extremely differentfrom those at normal mid-ocean ridges,whichhavelittleornosedimentarycover.Ataspreadingridge,MORBmagmasareextruded onto the sea floor, where theyquench to form pillow lavas and/or hyalo-clastite.Accompanyingmetamorphismisofahydrothermal-relatedocean-floortype.Incontrast,duringridge-trenchcollision,nor-mal-MORBmagmasinvadethebaseoftheaccretionaryprismandarerarelyextrudedontotheseafloor.Layeredgabbros,high-
temperature metamorphic rocks and gra-nitic rocks, and hybrid igneous rocks areformedbyfractionalcrystallizationofnor-malMORB,assimilationoftheaccretionaryprism, metamorphism, and anatexis, andmixingofmantle-derivedmaficmagmaandcrust-derivedanatecticmelts.Thus,juvenilecontinentalcrustisformedinfore-arcset-tingsthroughtheseintegratedprocesses,inmarkedcontrasttotheprocessesatnormalmid-oceanridges.AsnotedbyNelsonandForsythe(1989),ridgesubductionisoneofthemostimportanteventsforgenerationofcontinentalcrustthroughtheinteractionofasthenospheric mantle and supracrustalrocksaftertheonsetofplatetectonics.
ACKNOWLEDGMENTS
SupportedinpartbytheMinistryofEducation,Science,
and Culture of Japan grants 05640524 and 05231105. WethankPaulMorris,VirginiaSisson,andJamesHibbardforconstructivereviewsandsuggestions,andYasuoIkeda,KenjiTomura,andYujiOrihashiforhelpwithanalyticalwork.
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ManuscriptreceivedMay31,1995
RevisedmanuscriptreceivedSeptember11,1995ManuscriptacceptedSeptember25,1995
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The Island Arc
(2003)
12,
190–206
Blackwell Science, LtdOxford, UK
IARThe Island Arc1038-48712003 Blackwell Science Asia
122June 2003
390
Paleozoic ophiolites and blueschists
A. Ishiwatari and T . Tsujimori
10.1046/j.1038-4871.2003.00390.x
Original Article190206BEES SGML
*Correspondence.
Received 13 May 2002; accepted for publication 8 January 2003.
© 2003 Blackwell Publishing Asia Pty Ltd.
Thematic Article
Paleozoic ophiolites and blueschists in Japan and Russian Primorye in the
tectonic framework of East Asia: A synthesis
A
KIRA
I
SHIWATARI
1,
*
AND
T
ATSUKI
T
SUJIMORI
2
1
Department of Earth Sciences, Faculty of Science, Kanazawa University, Kanazawa 920-1192, Japan (email:
geoishw@kenroku.kanazawa-u.ac.jp) and
2
Research Institute of Natural Sciences, Okayama University of Science,
Okayama 700-0005, Japan
Abstract
Ophiolites and high-pressure (HP) metamorphic rocks are studied to test
continuation of Paleozoic and early Mesozoic geological units from Japan to Primorye overthe Japan Sea. The early Paleozoic ophiolites are present on both sides, and the latePaleozoic ophiolite of south-western Japan may also have its counterpart in Primorye. The
Shaiginskiy HP schist and the associated Avdakimov gneiss in Primorye, both tectonicallyunderlying the early Paleozoic ophiolitic complex, yield a 250-Ma phengite and hornblendeK–Ar age, which is intermediate between those of the Renge (280–330 Ma) and Suo (170–220 Ma) blueschists in south-western Japan. This age also coincides with that of the coesite-bearing eclogites in the Sulu–Dabie suture in China and several medium-pressuremetamorphic rocks in East Asia. On the basis of these results and other geological data, theauthors propose the ‘Y aeyama promontory’ model for an eastward extension of the Sulu–Dabie suture. The collision suture warps southward into the Y ellow Sea and detours aroundKorea, turns to the north at Ishigaki Island in the Y aeyama Archipelago of Ryukyu, where
it changes into a subduction zone and further continues toward south-western Japan andPrimorye. Most ophiolites from this area represent crust–mantle fragments of an islandarc–back-arc basin system, and the repeated formation of ophiolite–blueschist associationsmay be due to the repetition of the Mariana-type non-accreting subduction and Nankai-type accreting subduction.
Key words:
Japan Sea, Khanka terrane, Korea, Sikhote Alin, Sulu–Dabie suture, Y aeyama
promontory.
INTRODUCTION
More than a half century ago, long before theestablishment of plate tectonics, Kobayashi (1951)proposed a rifting–drifting hypothesis for the ori-gin of the Japan Sea. From a geological point ofview, the rifting–drifting theory requires theoccurrence of equivalent pre-Tertiary geologicalunits on both sides of the Japan Sea. For example,the Appalachian and Caledonian belts on bothsides of the Atlantic Ocean, respectively , representseparated fragments of a single early Paleozoicorogenic belt, and include Early Ordovician ophi-
olites of identical age (Dunning & Pedersen 1988).
Recent paleomagnetic results indicate a fast
drifting of Japan at
ca
15 Ma in the manner of the
opening of a pair of hinged doors (Otofuji &Matsuda 1984; Otofuji
et al
. 1985). Some basalt
samples drilled from the Japan Sea floor have
40
Ar-
39
Ar ages of 15–25 Ma (Kaneoka
et al
. 1992). These
data coupled with other geophysical and geochem-ical data result in various models for the Mioceneback-arc opening process (Nohda
et al
. 1988;
Tamaki & Honza 1991). However, evidence and
consideration for an original geological continuitybetween Japan and Russian Primorye are notconclusive.
The Japanese Islands are mainly composed of
accretionary complexes of Paleozoic, Mesozoic and
Paleozoic ophiolites and blueschists
191
Cenozoic ages. Every accretionary complex is
characterized by the ‘oceanic plate stratigraphy’,which is composed of fragments of oceanic crustand seamounts, chert and/or pelagic limestone, sil-iceous shale, sandstone and conglomerate (or olis-tostrome) in younging order (Isozaki 1996).Kojima (1989) pointed out that the Jurassic accre-
tionary complexes showing the same age–lithologyrelationship are found on both sides of the JapanSea, namely in south-western Japan (Mino-Tambabelt) and in the Sikhote Alin terrane in Primorye(Samarka zone) and the adjacent Chinese terri-tory (Nadanhada zone). However, detailed com-parison of Paleozoic ophiolites and blueschists onboth sides of the Japan Sea has not previouslybeen attempted.
The present paper reports petrologic and geo-
chronologic similarity of the Paleozoic igneous andmetamorphic rocks on both sides of the Japan Seaon the basis of our recent cooperative works withRussian geologists, and discusses configuration of
the late Paleozoic–early Mesozoic collision sutureand geotectonic significance of the multiple ophio-lite–blueschist assemblages in East Asia. Numericage data and detailed petrology of the dated sam-ples will appear elsewhere.
PALEOZOIC OPHIOLITES AND BLUESCHISTSIN JAPAN
SOUTH-WESTERN JAPAN
The Japanese Islands bear ophiolitic complexes ofvarious ages ranging from early Paleozoic to Cen-ozoic, forming a Phanerozoic multiple ophiolitebelt (Ishiwatari 1991, 1994). Paleozoic ophiolites(Fig. 1) occupy a higher structural position in thepiling nappes of the accretionary complexes. Insouth-western Japan (Fig. 2) the Oeyama ophioliteof Cambro-Ordovician age occupies the higheststructural position, and tectonically overlies thelate Paleozoic Renge blueschist, the Y akuno ophi-olite, and the Permian Akiyoshi (and Ultra-Tamba)accretionary complexes. These tectonic units arebounded by thrust faults (Ishiwatari
et al
. 1999).
The Oeyama ophiolite is mainly composed of
residual peridotite with podiform chromite depos-its and minor gabbroic rocks; basaltic volcanicrocks are completely absent. The peridotite is lher-zolitic in the eastern part (Oeyama body , spinel Cr#(Cr/(Al
+
Cr))
=
0.3; Kurokawa 1985), but is
harzburgitic in the western part (Tari-Misaka body ,spinel Cr#
=
0.5; Arai 1980). The peridotite includes
metagabbro and amphibolite bodies, which have
Fig. 1
Distribution of Paleozoic ophiolites
and Paleozoic–early Mesozoic blueschists inJapan.
192
A. Ishiwatari and T. Tsujimori
hornblende K–Ar ages of 444–464 Ma (Nishimura
& Shibata 1989). The kyanite- and staurolite-bearing HP metagabbro (troctolitic cumulate ori-gin) of higher metamorphic pressure in the Oeyamaperidotite body also has a similar hornblende K–Arage of 403–443 Ma (Tsujimori 1999; Tsujimori
et al
.
2000c; Tsujimori & Ishiwatari 2002).
The Y akuno ophiolite consists of a relatively
complete succession composed of harzburgite tec-tonite (spinel Cr#
=
0.6–0.8), dunite–wehrlite–
clinopyroxenite cumulate, metagabbro, amphibo-lite and metabasalt with abundant black shale(Ishiwatari 1985a). Middle–Late Permian radiolar-ian fossils were identified from the black shaleintercalated with basalt lavas (Kurimoto & Makim-oto 1990). Hornblende K–Ar dates of metagabbrorange from 241
±
12 to 278
±
10 Ma (Shibata
et al
.
1977), and conventional zircon U–Pb ages of theplagiogranite are 282
±
2 and 285
±
2Ma (Herzig
et al
. 1997), indicating an Early Permian igneous
age and slightly later metamorphism for this ophi-olite. In contrast, Sano (1992) reported a 421
±
54-
Ma Nd–Sm whole-rock isochron age (8 points) forthe metagabbro and plagiogranite, and a 311
±
65-
Ma Nd–Sm rock–clinopyroxene isochron age forthe metabasalt, suggesting a polygenetic naturefor the Y akuno ophiolite. However, coincidence ofthe Nd–Sm age of the Y akuno metagabbro and theK–Ar age of the Oeyama metagabbro may notindicate their cogenetic relationship. They shouldbe compared with the age data for the samemethod. Herzig
et al
. (1997) point out that the iso-
chron from Sano (1992) might be a mixing line.
The basaltic and gabbroic rocks of the Y akuno
ophiolite show transitional mid-ocean ridge basalt(T-MORB) chemistry in the eastern part and arc–tholeiite chemistry in the western part (Ishiwatari
et al
. 1990a). The metamorphic grade increases
from the prehnite–pumpellyite facies in the upperbasalt section to the granulite facies in the lowergabbro and ultramafic section. The granulite-facies metacumulate at the Moho is a gneissose
Fig. 2
Geological units in the main part of the Inner Zone of south-western Japan. Kyoto–Tottori cross-section is shown in the inset. ISTL, Itoigawa–
Shizuoka Tectonic Line; MTL, Median Tectonic Line; MB, metamorphic belt; AC, accretionary complex. Unit boundaries are mainly a fter Ishiwatari (1994),
Tsujimori (1998), and Geological Survey of Japan (1992).
Paleozoic ophiolites and blueschists
193
metagabbro composed of aluminous two
pyroxenes, plagioclase, and aluminous spinel(Ishiwatari 1985b). Such an aluminous-spinel met-agabbro is rare among ophiolitic complexes, andhas so far been reported only from Bikin, Pri-morye (Vysotskiy 1994) and Tonsina, Alaska(DeBari & Coleman 1989), although analogousrocks are known from meta-igneous complexes oflower continental crust (Rivalenti
et al
. 1981) and
lower crustal xenoliths in island arc (Francis 1976)and oceanic plateau (Grégoire
et al
. 1998).
The Renge blueschist occurs either as thin tec-
tonic slices or blocks in serpentinite mélangesunderlying the Oeyama ophiolite (Tsujimori 1998),and has a phengite K–Ar age of 320 Ma (Tsujimori& Itaya 1999; Fig. 3). The metamorphic assemblageof the mafic schist ranges from lawsonite blueschistto epidote blueschist and further into eclogitic rocksin the Omi area (Tsujimori
et al
. 2000a,b; Tsujimori
2002). The Renge metamorphic belt also includessome relatively low-pressure, high-temperaturemetamorphic rocks including oligoclase–biotiteschist and amphibolite. The Joetsu metamorphicbelt also bears typical blueschist and relativelyhigh-temperature pelitic schists with phengite K–Ar ages of 308 and 289 Ma (Y okoyama 1992), and isthought to be an eastern extension of the Renge belt(Fig. 1).
The Suo metamorphic belt also consists of high-
pressure (HP) metamorphic rocks includingpumpellyite–actinolite schist and epidote blue-schist (mostly winchite schist; Nishimura 1998).Some pelitic rocks bear lawsonite (Hayasaka 1987;Watanabe
et al
. 1987, 1989) but lawsonite–
glaucophane and pumpellyite–glaucophane assem-blages are absent. Phengite K–Ar ages are220
±
7Ma in the Nishiki area and 170–190 Ma in the
eastern areas (Nishimura 1998; Fig. 3). The schistswith the same K–Ar age are also known from theKurosegawa belt of the Outer Zone of south-west-
ern Japan and Ishigaki Island of Ryukyu (Fig. 3).
NORTH-EASTERN JAPAN
Early Paleozoic Miyamori ophiolite forms thebasement of a nearly complete Paleozoic–Mesozoicsedimentary sequence of the South Kitakami beltranging from Ordovician to Jurassic in age(Tazawa 1988, 2000). The Miyamori ophiolite con-sists mostly of wehrlitic cumulate and depletedmantle harzburgite (spinel Cr#
=
0.4–0.8) with
hornblende and phlogopite, which is interpreted tohave been upper mantle of an island arc (Ozawa1988). Minor lherzolite patches (spinel Cr#
=
0.1–0.4) of kilometric sizes are included in the depleted
harzburgite, and are interpreted to have been arelict source mantle that originated in a back-arcbasin (Ozawa 1988). Hornblende K–Ar ages of fourgabbroic rocks cutting peridotite of this ophiolite
Fig. 3
Isotopic ages of ophiolitic and metamorphic rocks in south-
western Japan, Russian Primorye and related areas. (a) K–Ar ages ofhigh-pressure schists in Primorye and Japan. Data sources: Kovalenkoand Khanchuk (1991); Tsujimori and Itaya (1999); Nishimura (1998) andthe references therein. Data for Primorye are mostly based on our unpub-lished results.
194
A. Ishiwatari and T. Tsujimori
range from 445 to 485 Ma, although other horn-
blendite (421 Ma) and amphibolite (369 Ma) yieldyounger ages (Ozawa
et al
. 1988). These data indi-
cate an Ordovician or earlier age for the Miyamoriophiolite, the same age as the Oeyama ophiolite insouth-western Japan.
The Motai blueschist belt (Maekawa 1988) in the
South Kitakami belt also bears pyroxene amphib-olite blocks with hornblende K–Ar ages of 479
±
24
and 524
±
26 Ma (Kanisawa
et al
. 1992). The K–Ar
age of a calk-alkaline tonalite dike cutting the meta-morphic rocks is 457 Ma (Sasada
et al
. 1992). A
monazite chemical Th-U-total Pb isochron method(CHIME, or electron microprobe method) age of430
±
10 Ma is also reported from the paragneiss in
the 350-Ma granitic complex (Suzuki & Adachi1991). These data indicate that ophiolites and meta-morphic and granitic rocks comprised an early
Paleozoic active continental margin or mature
island arc. The 300-Ma muscovite K–Ar age(Kawano & Ueda 1965) and 225
±
11- and 239
±
12-
Ma hornblende K–Ar ages (Kanisawa
et al
. 1992) of
garnet–epidote amphibolites from the Y amagamiarea suggest that the blueschist metamorphism inthe southern part of the Motai belt is contemporarywith the Renge and Suo blueschists of south-western Japan. However, late Paleozoic ophiolite(Y akuno) and accretionary complexes (Akiyoshiand Ultra-Tamba) are absent in north-easternJapan. The early Paleozoic ophiolitic–granitic base-ment and the sedimentary cover of Silurian–Jurassic ages in the South Kitakami belt thrustover the Jurassic accretionary complex of theNorth Kitakami belt (Tazawa 1988, 2000).
Fig. 3
(
Continued
)
(b) K–Ar ages for ophiolitic complexes in Primorye and Japan. Data sources: Khanchuk
et al
. (1996); Kovalenko and Khanchuk
(1991); Nishimura and Shibata (1989); Shibata
et al
. (1977); Sano (1992); Ozawa
et al
. (1988). K–Ar data for Primorye are mostly based on our unpublished
results. (c) Muscovite–phengite K–Ar ages of schists in the Kurosegawa belt, Outer Zone of south-western Japan (compilation in Tsujimori
et al
. 2000c).
(d) Nd–-Sm ages and zircon U–Pb ages of the ultrahigh-pressure metamorphic rocks in the Sulu and Dabie areas, China (after Ames
et al
. 1993; Li
et al
.
1993). ( ), hornblende K–Ar age; ( ) muscovite–phengite K–Ar age; ( ), muscovite–phengite Ar–Ar age; ( ), zircon U–Pb age; ( ), Nd–S m min-
eral isochron; ( ), Nd–Sm rock isochron.
Paleozoic ophiolites and blueschists
195
The radiometric ages of pre-Cretaceous ophi-
olitic and metamorphic rocks in south-westernJapan are summarized in Fig. 3. The ophiolite agesare centered on two peaks at approximately 450 Ma(Oeyama ophiolite) and 280 Ma (Y akuno ophiolite).North-eastern Japan bears the 450-Ma ophiolitesbut lacks the 280-Ma ophiolites. The HP metamor-phic rocks are concentrated around two otherpeaks at approximately 300 Ma (Renge blueschist)and 200 Ma (Suo blueschist), although minor olderHP metamorphic rocks of 400–450 Ma are alsopresent in association with the Oeyama ophioliteand the Kurosegawa mélange in Shikoku Island.
PALEOZOIC OPHIOLITES AND BLUESCHISTS IN PRIMORYE, RUSSIA
The Primorye territory of Russia is geologicallydivided into two parts: the Khanka terrane and theSikhote-Alin terrane (Fig. 4). The Khanka terranemay be a part of a larger continental block includ-ing the Bureya and Jiamusi blocks to the north
(Khanchuk 2001), composed of Precambrian conti-nental basement covered by thick Cambrian cal-careous sediments and post-Silurian continentalsediments. The Sikhote-Alin terrane mainly con-sists of Mesozoic accretionary complexes of green-stone, chert, limestone, shale and sandstone, whichare intruded by Cretaceous granites and coveredby Cretaceous–Tertiary volcanics.
KHANKA OPHIOLITE
The Khanka terrane bears some ophiolitic bodies inthe west-north-west-trending Spassk zone, which isalmost perpendicular to the general trend of theSikhote-Alin terrane. Shcheka
et al
. (2001) describe
an ophiolitic sequence of serpentinite (harzburgite),pyroxenite, gabbro and basalt that are emplacedonto the Early Cambrian fossiliferous limestone–shale formation and covered by Middle Cambrianconglomerate including abundant detrital chro-mian spinel. The chromian spinel grains from ser-
Fig. 4
Geological units of Russian Primorye
(simplified after Khanchuk
et al
. 1996; with addi-
tion of Shaiginskiy blueschist).
196
A. Ishiwatari and T. Tsujimori
pentinite and conglomerate are high in Cr# (0.50–
0.75) and very low in TiO
2
(
<
0.2 wt%) as those in
other ophiolites. Although the area is free fromregional metamorphism, ferrichromit rims aredeveloped in most chromian spinel grains; some areunusually rich in MnO (up to 19 wt%), suggestingocean-floor hydrothermal alteration prior to theemplacement. Shcheka
et al
. (2001) conclude that
the Khanka ophiolite may represent oceanic litho-sphere formed in a small rift zone and immediatelyemplaced onto the adjacent passive continentalmargin prior to the circum-Pacific orogeny .
SIKHOTE ALIN OPHIOLITES
Major ophiolitic complexes of the Sikhote Alin ter-rane are aligned in a north-north-east directionparallel to the eastern margin of the Khanka ter-rane. The trend also parallels the Central SikhoteAlin Fault. The ophiolitic complexes are distrib-uted in the Sergeevka, Kalinovka and Bikin areasfrom south to north.
The Sergeevka metagabbro body forms a north-
east-trending massif of 30
¥
130 km, the largest
mafic body in Primorye. Gneissose hornblendemetagabbro occupies more than 80% of the area inassociation with some granitic and troctolitic intru-sions, as well as various metamorphic rocks suchas gneiss, amphibolite and marble. Conventionalzircon U–Pb ages of 528
±
3Ma are reported for
gneissose metagabbro, 504
±
3Ma for gneissose
diorite, and 493
±
12 Ma for granite (Khanchuk
et al
. 1996). Mishkin
et al
. (1970) reported a mus-
covite K–Ar age of 529 Ma for another granite anda hornblende K–Ar age of 622 Ma for garnetamphibolite, but Tsujimori (unpubl. data) obtainedeight hornblende K–Ar ages of metagabbroswithin a narrow range between 430 and 470 Ma(Fig. 3). These hornblende K–Ar ages are similarto those of the Oeyama ophiolite in south-westernJapan and the Miyamori ophiolite in north-easternJapan (Fig. 3). Khanchuk
et al
. (1996) consider
that this body forms a part of the continental mar-gin of the Khanka block, hence it is not an ophio-lite. However, the dominantly mafic nature of thisbody , its occurrence as a nappe overlying youngerblueschist and accretionary complex, and its posi-tion located on the same line as the other Sikhote-Alin ophiolites (Fig. 4) suggest that the Sergeevkamassif is a dismembered ophiolite body .
The Kalinovka ophiolite group is composed of
three north-east-trending ophiolitic bodies ofapproximately 5
¥
40 km having an
en echelon
arrangement. These bodies are composed of dun-ite, troctolite, wehrlite, clinopyroxenite, olivine
gabbro, hornblende gabbro, plagiogranite, pillowbasalt and minor amphibolite and granite. Thechert and limestone associated with the pillow lavabear conodont fossils of Late Devonian–EarlyPermian ages (Vysotskiy 1994). The K–Ar age of
410
±
9Ma is the only reported age determined for
very K-poor hornblende (Kemkin & Khanchuk1994). Our preliminary hornblende K–Ar datingfor the metagabbro at Medvezhy Kut nearBreyevka yields 230 Ma, which is younger thanthat of the Y akuno ophiolite (Fig. 3). Vysotskiy(1994) describes an olivine–plagioclase reactionrelationship to form aluminous spinel–pyroxenesymplectite in troctolite, and Khanchuk andPanchenko (1991) report garnet metagabbro. The
tectonic superposition of the Kalinovka ophioliteover the Jurassic accretionary complex of theSamarka zone with an intervening older accretion-ary complex called the Udeka zone resembles ananalogous relationship in south-western Japan,where Y akuno ophiolite thrust over the JurassicTamba zone with the Permian Ultra-Tamba zone in
between (Kojima
et al
. 2000).
The Bikin ophiolite group is composed of three,
north–south-trending ophiolitic bodies of 1
¥
2–
3km in size; namely the Oronsky , Zalominsky and
Soldinsky bodies (Vysotskiy 1994; Vysotskiy
et al
.
1995). Dunite, harzburgite, wehrlite, orthopyrox-enite, and aluminous spinel-bearing noritic gabbroare associated with Late Permian basaltic pillowlava and siliceous volcanic rocks as well as a ser-pentinite mélange. The aluminous spinel-bearing,olivine-free gabbro with very aluminous pyroxenes(Al
2
O
3
>
8wt%) at the Moho is evidence for unusu-
ally thick oceanic crust (Ishiwatari 1985a), and itsoccurrences on both sides of the Japan Sea (in theYakuno and Bikin ophiolites) suggest original con-
tiguity of the Paleozoic ophiolite belt.
SIKHOTE ALIN BLUESCHIST
The Shaiginsky blueschist occurs as windows andthin thrust sheets beneath the Sergeevka ophioliticbody . The epidote blueschist bears crossite andbarroisite. Pelitic rocks are of a higher grade thanthe garnet zone, and some samples bear oligoclase(An
18
), although biotite is completely absent. Gar-
net preserves progressive normal zoning withdecreasing Mn toward the rim; some showsreverse zoning at the rim. Piemontite-bearing sil-iceous schist is also present near Partisansk. TheShaiginsky schists yield phengite K–Ar ages of230–250 Ma (Fig. 3). These age data lie between
Paleozoic ophiolites and blueschists
197
those of the two HP metamorphic belts in south-
western Japan (i.e. the Renge (280–330 Ma) andSuo belts (170–220 Ma)). The Shaiginsky blueschistis associated with the ‘ Avdakimov gneiss’ composedmainly of hornblende gneiss with marble andcoarse-grained garnet amphibolite. Although aPrecambrian Rb–Sr mineral whole-rock isochronage was reported from this gneiss complex, ourhornblende K–Ar ages are centered at 250 Ma(Fig. 3), the same age as the Shaiginsky blueschistcomplex. Kovalenko and Khanchuk (1991) reporteda 255-Ma and 290-Ma phengite K–Ar age for peliticschists of the Shaiginsky complex; our K–Ar dataare centered at 250 Ma (Fig. 3). These ages areintermediate between those of the Renge (280–330 Ma) and Suo (170–220 Ma) blueschists of south-western Japan. Nevertheless, the K–Ar age of theSuo metamorphic rocks varies significantly fromarea to area (Fig. 3); some metamorphic rocks inwestern Kyushu and in the Kurosegawa Belt havethe same age as the Shaiginskiy blueschist.
GEOLOGICAL CONTINUATION FROM JAPAN TO PRIMORYE
Even if we accept the rifting–drifting hypothesis,the Pre-Japan Sea configuration of the JapaneseIslands is not easy to restore. Some authors
assume that south-western Japan was locateddirectly to the south of Primorye and on the east ofthe Korean Peninsula (with Y amato Bank inbetween) as shown in Fig. 5 (Kojima 1989;Khanchuk 2001). In contrast, the South KitakamiBelt and associated accretionary complexes ofnorth-eastern Japan should already have beenplaced alongside Primorye or between Primoryeand south-western Japan in the Early Tertiaryprior to the opening of the Japan Sea. However,paleontologic data of Paleozoic and Mesozoic for-mations in Japan indicate that the South KitakamiBelt and the Kurosegawa Belt (the Outer Zone ofsouth-western Japan) were placed in the Chinesecontinental margin to the south of south-westernJapan in Cretaceous and earlier time (Otoh &Sasaki 1998; Tazawa 2000), and displaced north-ward through fast and extensive left-lateralstrike–slip movement.
Arakawa
et al
. (2000) mention the possibility
that the Hida belt does not belong to the Sino-Korean block, but has evolved as a part of the East-
Central Asian Orogenic Belt, which is a wide accre-tionary belt extending from Primorye to CentralAsia via north-eastern China and Mongolia along
Fig. 5
(a) Continuation of geological units from south-western Japan to Russian Primorye before opening of the Japan Sea. Position of south-western
Japan follows that of Kojima (1989). (b) Major geological events in Japan and Primorye.
198
A. Ishiwatari and T. Tsujimori
the northern margin of the Sino-Korean block.
Khanchuk (2001) suggests that the Laoelin–Grode-kovo belt and Cheongjin belt in the Russia–China–North Korea border area are possible extension ofthe ophiolite belts in south-western Japan.
PALEOZOIC-EARLY MESOZOIC BLUESCHISTS IN JAPAN AND PRIMORYE: RELATION TO THE COLLISION BELT IN CHINA
As noted in the previous section, some late Paleo-zoic–early Mesozoic blueschist facies rocks of
ca
250 (
±
100) Ma occur in southern Primorye (Shai-
ginsky blueschist: Kovalenko & Khanchuk 1991and the present study), northern Honshu (Motaibelt: Maekawa 1988; Kanisawa
et al
. 1992), west-
ern Honshu and Kyushu (Renge and Suo belts:Nishimura 1998; Tsujimori & Itaya 1999), andIshigaki Island of Ryukyu (Tomuru Formation:Nishimura
et al
. 1983; Faure
et al
. 1988). They are
mostly epidote blueschist in Primorye, Motai, Suoand Ishigaki, but typical lawsonite blueschist andeclogitic rocks occur in the Renge belt (Tsujimori1998; Tsujimori
et al
. 2000a,b). It should be noted
that HP metamorphism in Japan took place notonly in late Paleozoic–Early Mesozoic time but alsoin early Paleozoic time (Kurosegawa: Maruyama &Ueda 1974; Oeyama: Tsujimori 1999; Tsujimori
et al . 2000c) and in Cretaceous time (Sambagawa,
Nagasaki, and Kamuikotan). Thus, subduction-zone metamorphism repeatedly took place inJapan; a subduction zone also existed in Japan atca250 Ma, when collision of the Sino-Korean and
Yangtze cratons took place (Fig. 3).
DOES THE CHINESE COLLISION SUTURE GO TO KOREA?
The Chinese Dabie-Sulu ultrahigh-pressure
(UHP) metamorphic belt is believed to be a colli-sional suture between the Sino-Korean andYangtze blocks, which amalgamated during the
200–250-Ma period, on the basis of the Nd–Sm andU–Pb ages of the UHP metamorphic rocks (Ameset al . 1993; Li et al . 1993; Hacker et al . 1998; Jahn
1998). The presence of Triassic flysch and Jurassicmolasse along the suture also supports earlyMesozoic collision of the continental blocks (Li1996). However, K–Ar ages of phengite and horn-blende in the UHP metamorphic rocks are scat-tered widely , from the Proterozoic to the Mesozoic(Ishiwatari et al . 1990b; Li et al . 1994), possibly
due to excess argon inherited from their Precam-brian protoliths (Giorgis et al . 2000). The Chinesesuture is postulated to extend into the Korean Pen-
insula, namely into the Imjingang or Ogcheon belt(Ernst & Liou 1995; Ree et al . 1996), and further
continuing into the Hida (marginal) belt in Japan(Isozaki 1996, 1997). The latter idea is based onHiroi’s works of 1981 and 1983 (Hiroi 1981, Hiroi1983), which first correlated the Unazuki meta-morphic rocks in the Hida belt to the Ogcheon andImjingang (Y onchon or Y eoncheon) metamorphicbelts. In their model, the northern part of theKorean Peninsula belongs to the Sino-Korean
block, whereas its southern part belongs to theYangtze block. Kim et al . (2000) reported a Rb–Sr
mineral isochron age of 226 ±1.2 Ma for the mylo-
nite in the Gyeonggi (Kyonggi) massif on the southof the Imjingang belt, and interpreted it to repre-sent post-collisional, extensional ductile shear. Leeet al . (2000) correlated the early Proterozoic gran-
ulites of the Gyeonggi massif to that of the Y angtzecraton on the basis of zircon–monazite sensitivehigh mass-resolution ion microprobe (SHRIMP)U–Pb age and an Nd–Sm isotope model age (T
DM),
although they admit ‘it is probably not warrantedto attempt any tectonic correlation solely based onthe resemblance in T
DM ages or SHRIMP zircon
ages’.
The Imjingang metamorphic belt is a typical
Barrovian kyanite–sillimanite-type metamorphicbelt (Y amaguchi 1951). Ree et al . (1996) reported a
Nd–Sm mineral isochron age of 249 ±31 Ma, and
0.8–1.3 GPa and 630–790 ∞C metamorphic condi-
tions for the adjacent garnet amphibolite unit onthe south, and correlated it to the Permo-Triassicsuture in China, regarding the garnet amphiboliteas retrograded from UHP eclogite. However, ‘crit-ical evidence of UHP metamorphism such as eclog-ite, diamond and coesite remains to be found’ (Reeet al . 1996). Min and Cho (1998) identified a three-
stage metamorphic evolution of the Ogcheon belt:(1) Siluro-Devonian medium-pressure (0.5–0.8 GPaand 520–590 ∞C) metamorphism; (2) Triassic
regional retrograde metamorphism (0.1–0.3 GPaand 350–500 ∞C); and (3) Jurassic–Cretaceous ther-
mal metamorphism around granitoids. This meta-morphic history coincides with the structuraldevelopment of the Ogcheon belt as an early Pale-ozoic intracontinental rift zone evolved into anintracontinental fold–thrust belt without ophiolite(Cluzel et al . 1990), but is not consistent with the
Permo-Triassic intercontinental collision process
that involves HP metamorphism (Ernst & Liou1995). Moreover, another medium-pressure meta-morphic belt with a muscovite
40Ar–39Ar age of 200–
230 Ma is reported from the Fangshan area in the
Paleozoic ophiolites and blueschists 199
Western Hills of Beijing (Fig. 6), that is, in the mid-
dle of the Sino-Korean craton (Wang & Chen 1996).As aforementioned, there is no direct evidence ofUHP–HP metamorphism in the Korean metamor-phic belts; it is likely that the 200–250-Ma medium-pressure metamorphic belts such as Imjingang,Ogcheon and Fangshan took place in the intracon-tinental fold–thrust belts of the Sino-Korean cratonduring the Y angtze–Sino-Korean collision.
The most convincing criteria with which to
define tectonic affiliation of an area may be stratig-raphy of the sediments covering continental base-ment. The Paleozoic system of the Ogcheon zoneshows typical ‘Sino-Korean’ stratigraphy charac-terized by thick Cambro-Ordovician limestone,late Paleozoic coal-bearing sediments, and ‘thegreat hiatus’ in between (Fig. 7). This is differentfrom the Y angtze stratigraphy with thick Siluro-Devonian shale (Fig. 7). This suggests that nomajor suture exists between the northern andsouthern parts of Korea. Lee et al . (1998) also state
that the Korean Peninsula as a whole belongs tothe Sino-Korean craton at least from the late Pro-terozoic in view of the overall similarity in age,geology , petrography and geochemistry .YAEYAMA PROMONTORY HYPOTHESIS
In view of the eastward-convex, winding geological
structure over the Korean Peninsula, Teraoka et al .
(1998) proposed that the Sulu suture does notextend to Korea, but turns southward beneath theYellow Sea. They did not specify , however, where
the destination of the redirected suture is. Theirintensive studies on the chemistry of clastic gar-nets in the Japanese Cretaceous–Tertiary sedi-ments indicate that these eclogitic, pyrope (Mg)-rich and spessartine (Mn)-poor garnets occur insandstones of the Shimanto accretionary complexin the outer zone of the South-west Japan andRyukyu arcs (Takeuchi 1992; Teraoka et al . 1999),
whereas such garnets are not found from the Cre-taceous–Paleogene fore-arc and intra-arc basins ofsouth-western Japan. This finding suggests thatthe Sulu UHP belt does not extend to Korea andnorth-eastern China, but turns to the south into theprovenance area of the Shimanto sediments. Wepropose that the Sulu suture reappears at IshigakiIsland, Y aeyama Archipelago, southern Ryukyu,and continues into Japan and Primorye, detouringaround Korea (Fig. 6). Along this highly sinuous
Fig. 6 Proposed sinuous configuration of the
eastern elongation of the Sulu–Dabie suture ( ca
250 Ma) of China passing subduction zones ofRyukyu, south-western Japan and Russian Pri-morye but detouring around Korea. A prelimi-nary version of this diagram appeared inIshiwatari and Tsujimori (2001).
200 A. Ishiwatari and T. Tsujimori
suture line, continental collision took place in the
Chinese segment, whereas subduction of oceaniclithosphere took place in the Japanese–Russiansegment. This agrees with the argument of Ernstand Liou (1995) that UHP metamorphism takesplace in the continental collision segment, whilenormal HP metamorphism takes place in the oce-anic subduction segment, along a single suture line.
Such a sinuous configuration of the collisional
suture is also observed in the Alpine chain in theMediterranean area. It is noteworthy that the coes-ite-bearing UHP metamorphic rocks of the DoraMaira massif (Chopin 1984) occur at the north-western tip of the acute Adriatic (Apulian) prom-ontory , whose profile is clearly visible on thepresent-day seismic map (Mueller 1989). Ourmodel suggests that the Chinese UHP rocks occurat the southern tip of the Dabieshan promontory ofthe Sino-Korean Craton and at the northern tip ofthe Sulu promontory of the Y angtze craton. In thiscontext, the early Mesozoic HP schists of the Ish-igaki Island possibly represent the southern tip ofanother promontory of the Sino-Korean Craton,which we call Y aeyama promontory after theregional name for the southernmost RyukyuIslands. The late Paleozoic–early Mesozoic (200–
250-Ma) HP metamorphic belts in Ryukyu, Japanand Russian Primorye are suitable as an easternextension of the Chinese collisional suture of thesame age. The late Paleozoic–early Mesozoic(Indosinian) dextral ductile shearing reported fromthe Ogcheon belt (Cluzel et al . 1991; Otoh et al .
1999) is compatible with the reciprocal movementbetween the Sulu and Y aeyama promontories.
The Tananao schist complex of eastern Taiwan
has long been regarded as a late Paleozoic oro-genic belt (Fig. 4 of Cluzel 1991), but hornblendeK–Ar ages of the schist are younger than 90 Ma,and the associated gneiss and granite also give90 Ma or younger Rb–Sr and U–Pb ages (Yuliblueschist is as young as 10 Ma; Jahn et al . 1986).
This indicates that the late Paleozoic–early Meso-zoic metamorphic belt around the Y aeyama prom-ontory does not extend to Taiwan.
The Y aeyama promontory hypothesis provides
some insights for pre-Cretaceous paleogeographyof Japan. Paleozoic and Mesozoic fossil faunas indi-cate that the South Kitakami and KurosegawaBelts were situated further to the south of the Hidamarginal belt before the Late Cretaceous large-scale strike–slip movement (Otoh & Sasaki 1998;Tazawa 2000). These three belts show overall strati-
graphic similarity with the Khanka massif; theseterranes may have together developed along theactive continental margin of the Sino-Korean cra-ton, although the South Kitakami and KurosegawaBelts were later displaced toward the north bya Cretaceous left-lateral strike–slip movement(Tazawa 2000). Isozaki (1997) proposed thatJapanese accretionary complexes, including theSouth Kitakami block and ‘Y akuno oceanic plateau’,developed along the Y angtze continental margin,assuming that the Sino-Korean–Y angtze suturezone passes through central Korea and extends tothe Hida Mountains in south-western Japan. Asmentioned earlier, however, the fossil fauna andlithology of the Permo-Triassic cover of the Y akunoophiolite closely resemble those of Primorye(Nakazawa 1958). Permian strata of the Hida mar-ginal belt and South Kitakami belt also show a fau-nal kinship with those of north-eastern China andPrimorye (Tazawa 1993, 2000; Otoh & Y anai 1996;Otoh & Sasaki 1998). Our model infers that allJapanese Paleozoic terranes, except for accretedseamounts, have developed along the Sino-Koreanmargin. The South Kitakami and KurosegawaBelts may have been placed somewhere betweenKyushu and Ishigaki Island along the eastern mar-
gin of the Y aeyama promontory (Fig. 6).
Fig. 7 Schematic stratigraphic columns for the contrasting Paleozoic
sequences on the (a) Yangtze block (eastern Sichuan, lower Yangtze Val-ley) and (b) Sino-Korean block (Hebei, Shanxi; after Willis & Blackwelder1907). The paleozoic sequence in southern Korea is also of the Sino-Korean type.
Paleozoic ophiolites and blueschists 201
PALEOZOIC OPHIOLITES IN JAPAN AND
PRIMORYE: GEOTECTONIC IMPLICATIONS
Cluzel (1991) proposed that the Y akuno ophiolite
formed in a rift zone in the Sino-Korean continen-tal margin in Middle–Late Carboniferous time,and was emplaced onto the ‘Honshu block’ in themiddle Permian by the closure of the small seabasin between the rifted continental blocks. Such a‘Tethyan ophiolite’ model including continental
rifting, sea-floor spreading, and obduction may beapplicable to the Cambrian Khanka ophiolite inPrimorye, which may have formed before thebeginning of the circum-Pacific-type orogeny , butmay not apply for the other circum-Pacific ophio-lites (Ishiwatari 1994). The Y akuno ophiolite is tec-tonically underlain by the Permian Ultra-Tambaaccretionary complex, which is in turn underlainby the Jurassic Tamba accretionary complex. Eachof the two complexes consists of ‘oceanic platestratigraphy’ (Isozaki 1996); successive underplat-ing of oceanic and trench-fill sediments beneaththe Y akuno ophiolite may have developed theseaccretionary complexes. The Miyamori ophiolitealso thrust over the Jurassic accretionary complex(Tazawa 1988). These ophiolites did not thrust ontoold continental blocks as imagined in the Tehyanmodel. Instead, younger accretionary complexesformed beneath the old ophiolites.
In contrast, Isozaki (1996, 1997) proposed that
the Y akuno ophiolite with thick crust representsoceanic plateau, which has formed in the midst ofthe ocean by a superplume activity and lateraccreted to Japan. However, the sedimentarycover of the Y akuno ophiolite is thick black shalewith a restricted age of radiolarian fossils (middlePermian), which is incompatible with a long plate-
tectonic travel in the ocean.
Ishiwatari et al . (1990a) divided the gabbroic
rocks of the Y akuno ophiolite into two types: aMORB type in the eastern area and an island-arcbasalt (IAB) type in the western area, according tothe chemistry of coexisting clinopyroxene and pla-gioclase. They postulated the Y akuno ophiolite asbeing a cross-cut section of oceanic island arc andan adjacent back-arc basin, which was affected bya mantle plume. The granulite-facies metacumu-late represents the basal part of thickly developedmafic crust of the island arc and back-arc basin.The mantle section of the Miyamori ophiolite isalso interpreted as hydrous mantle beneath islandarc (Ozawa 1988). Mantle peridotite of the Oeyamaophiolite resembles that from the ocean floor, andbears some podiform chromitite with hydrous min-eral inclusions such as Na-phlogopite and par-
gasite (Matsumoto et al . 1995), suggesting either a
supra-subduction zone (SSZ) or a fast spreadingridge setting (Arai 1997). However, the SSZ set-ting is more preferable in view of the orbicularchromitite with a very high Cr# (0.76–0.85)reported by Y amane et al . (1988).
It should be noted that the ophiolite sequence is
actually exposed in the submarine trench wallsaside the Mariana arc (Bloomer & Hawkins 1983)and Tonga arc (Bloomer & Fisher 1987; Fig. 8). Itis important that back-arc spreading is activebehind these island arcs, and the accretionarycomplex is currently absent in the subductionzone. Moreover, lawsonite-bearing blueschistblocks were drilled from serpentinite seamounts(diapirs) in the Mariana fore-arc area (Maekawaet al . 1993, 1995; Fig. 8). The ophiolite–blueschist
association is well demonstrated in the Japan–Primorye area such as the Oeyama ophiolite–Renge blueschist, Y akuno ophiolite–Suoblueschist, Miyamori ophiolite–Motai blueschistand Sergeevka ophiolite-Shaiginskiy blueschist. Incontrast, subduction zones such as the JapanTrench and Nankai Trough have formed huge
accretionary complexes from the Cretaceous tothe present (Taira 1985; Fig. 8).
It is likely that the accretion period and non-
accretion period as represented by the presentNankai Trough and Mariana Trench, respectively ,have repeated one after another and from segmentto segment in the history of the Japan–Primoryeaccretionary orogenic belt. The ophiolite–blue-schist period with tectonic erosion at the subduc-tion zone may have been followed by a period ofmassive accretion. This idea is quite compatiblewith the geochemical island-arc and marginal-basin signatures of the ophiolitic rocks.
CONCLUSIONS
Age, lithology , and structural position of the Pale-
ozoic ophiolites and blueschist in south-westernJapan and Primorye strongly support original con-tinuation of the geological units over the two sidesbefore the Miocene opening of the Japan Sea. Thenewly obtained K–Ar ages of the blueschists andassociated gneiss complex of Primorye (250 Ma)coincide with the widespread metamorphic eventsin East Asia such as the UHP metamorphism of theSulu–Dabie suture and medium- and low-pressuremetamorphism in Japan (Unazuki and Hida), Korea(Imjingang and Ogcheon), and northern China
202 A. Ishiwatari and T. Tsujimori
(Fangshan). The wide-scale convergence resulted
in the UHP metamorphism along the continentalcollision zone, but ordinary HP metamorphism tookplace along the sinuous oceanic extension of thesame suture passing Ryukyu, Japan and Primoryebeyond the Y aeyama promontory , detouring aroundKorea. Although 250-Ma HP schists are rare in
Japan, slightly older (Renge) and younger (Suo)blueschists are preserved in many places, suggest-ing persistent subduction. The repeated formationof the ophiolite–blueschist assemblages and the tec-tonically underlying, younger accretionary com-plexes suggests repetition of the non-accretingsubduction as in the present Mariana Trench andthe accreting subduction as in the Nankai Troughthrough development of the orogenic belts in Japanand the Russian Far East.
ACKNOWLEDGEMENTS
We thank Professors A. I. Khanchuk, S. V .
Vysotskiy , S. A. Shcheka and S. V . Kovalenko of the
Far East Geological Institute of the Russian Acad-
emy of Sciences in Vladivostok for their coopera-tion in the field and sample preparation. ProfessorTetsumaru Itaya of Okayama University of Sci-
ence is thanked for providing equipment andinstructions for K–Ar dating for the second author.
Professors J. G. Liou and H. Maekawa arethanked for constructive reviewing of the manu-script. The first author acknowledges Grant-in-Aidfor Scientific Research (C)-(2)-10640462 and-14540447 provided by the Ministry of Education,Japan. The second author acknowledges Grant-in-Aid for JSPS Fellows.
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Journal of Asian Earth Sciences 72(2013) 88-101 ContentslistsavailableatSciVerseScienceDirect Journal of Asian Earth Sciences ELSEVIER journal homepage: www.elsevier.com/locate/jseaes Review Late Paleozoic-Mesozoic tectonic evolution of SW Japan: CrossMark A review - Reappraisal of the accretionary orogeny and revalidation of the collisional model Jacques Charvet Institut des Sciences de la Terre dOrléans (ISTO),Universite d'Orléans, ISTO,UMR7327,45071 Orleans,France ARTICLEINFO ABSTRACT Article history: This paper makes a review of the interpretations of the tectonic evolution of SW Japan during the last three Available online 17 May 2012 decades. In the late 1970s, the dominant model was the so-called “Pacific-type orogeny", emphasizing the purported absence of nappes and the contrast with the alpine chains, and interpreting the evolution as due Keywords: to a steady oceanic subduction since the Paleozoic time. In the 80s, the discovery of the actual structure Japanese Islands made of a pile of large thrust sheets led authors to propose collisional models, involving the intermittent Collisional orogen/accretionary orogen underthrusting of buoyant blocks like micro-continents. At the same time, the use of high-resolution bio- Late Paleozoic Mesozoic stratigraphy allowed several authors to recognize ancient accretionary wedges, with a reconstructed ocean plate stratigraphy of individual accreted units, especially in the Tanba and Shimanto zones. Also, precise Tectonics radiometric dating permitted the distinction of metamorphosed units, especially in Sanbagawa and Geodynamics Shimanto belts. As a result of these new data, since the 1990s, the plate tectonic interpretation of the ary processes linked to a steadily oceanic subduction, with an episodic ridge subduction: the so-called "Miyashiro-type orogeny". The review of different data leads to the following conclusions. The structure one encountered in collisional orogens. The geodynamic mechanisms advocated for the tectonic building oceanic subduction with the intermittent “collision" (actually subduction) of an active ridge or seamount chain is unable to build such structures, as this process induces in fact an acceleration of the tectonic ero- sion and collapse of the upper plate; the underthrusting of a micro-continent or mature arc is likely needed. rological and new geochemical data from the literature strongly support the existence, beneath the nappes of accretionary complexes, of continental bodies showing affinities with South China, from which they were once separated. The episodic collision, underthrusting, of such blocks was responsible for the tectonic piling. Tectonic erosion plaid likely a major role in removing material during the intervening subduction stages. A revised geodynamic model, implying the collision of the Honshu, South Kitakami-Kurosegawa, and Shimanto Blocks, is proposed for explaining the three orogenic crises which took place respectively at around 240, 130, and 80-60 Ma ago in SW Japan. The paleogeographic position and affinity of the Hida block with surrounding units, in the hinterland, are still unclear. More work is needed to solve this question. @ 2012 Elsevier Ltd. All rights reserved. Contents 1 Introduction .. 89 2. The discovery of nappes: invalidity of the Pacific-type orogeny, proposal of the collisional model. 89 3. Recent interpretations: recognition of large thrust sheets but return to the accretionary model with ridge subduction. 90 4. Invalidity of the accretionary orogeny: revalidation of the collisional model . 92 4.1. Tectonic considerations. 92 4.2. Geodynamic implications. 92 4.3. Exhumation of Sanbagawa HP metamorphic rocks. . 94 4.4. Petrological and geochemical evidence . 95 4.5. Summary ... . 95 E-mail address: jacques.charvet@univ-orleans.fr 1367-9120/$ - see front matter @ 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/jseaes.2012.04.023 J. Charvet/Journal of Asian Earth Sciences 72(2013） 88-101 89 5. Role of tectonic erosion.. . 95 6. Origin of colliding blocks. 96 7. Geodynamic model ...... 98 8. Conclusions. 98 Acknowledgements ... 99 References .. 99 1. Introduction and different tectonic events, I will restrict my point to SW Japan, to the SW of the Tanakura fault (Fig. 1) and mainly to the tectonic The Japanese Islands are often considered as a typical example development from Late Paleozoic to Paleocene. of accretionary orogen (e.g. Cawood et al., 2009), built during the wedges composed of material coming from the upper plate by ero- 2. The discovery of nappes: invalidity of the Pacific-type sion and mainly from the downgoing plate (accretion s. st.), includ- orogeny, proposal of the collisional model ing deep-sea sediments, oceanic plateaus, HP metamorphic rocks. - si d an n ss s on na As mentioned before, the theory of "Pacific-type orogeny" (Mat- suda and Uyeda, 1971) was dominant in the 1970s. This model melting of the lower plate crust and/or mantle wedge. The domi- emphasized the role of paired metamorphic belts (Miyashiro, nant model expressed today in the literature presents the Japanese 1961) and the assumed total absence of nappes in Japan and, due Islands as a segment of subduction-related orogen built along the to this lack, the strong contrast with the Alpine chains. Although western Pacific margin, since at least the Mesozoic, by successive the concept of paired metamorphic belts can be discussed and underthrusting of tectonic units corresponding to accretionary questioned (e.g. Brown, 1998, 2010), this point will not be ad- complexes, either sedimentary or affected by high P/T metamor- dressed in detail here and treated only incidentally; I will concen- phism (e.g. Isozaki, 1996, 1997; Maruyama, 1997; Isozaki et al., trate on the tectonic and geodynamic aspects. 2010). The episodicity of orogeny and production of granites is In the 1970s, the geometry of structures was poorly known and now assigned to the intermittent subduction of an oceanic ridge, the stack of nappes ignored, despite some previous works advocat- ing their existence (e.g. Fujimoto, 1937). However, in the late A similar concept of subduction-related orogeny, piling Permian 1970s and early 1980s, two developments induced a big change to Recent accretionary complexes, corresponds to the “Nippo- in the understanding of the geometry of the different units, leading nides" of Sengor and Natal'in (1996). to the recognition of the structural scheme admitted today (Fig. 2). Basically, this interpretation is quite similar to the theory of The first one was a revolution in the Japanese stratigraphy, due to "Pacific-type orogeny" (Matsuda and Uyeda, 1971), which was the use of conodont and radiolarian faunas. The second one was dominant in the 1970s. This model emphasized the role of paired the input of modern structural studies, by Japanese and foreign sci- metamorphic belts and insisted on one structural point: the as- entists, in particular the French team with an experience of alpine sumed total absence of nappes in Japan and, due to this lack, the belts. The dating by conodonts and radiolaria showed that some strong contrast with the Alpine chains. Therefore, in parallel with formations, previously considered as Permian in the Inner zones the development of the plate tectonic theory and the discovery e ss isnis i u e ue s jo of the great importance of subduction processes, an extrapolation of the present situation was made in the past; the Japanese Islands should be interpreted as tectonic windows (Hayasaka and Hara, were considered as a permanent island-arc and a giant accretion- 1982; Faure and Caridroit, 1983). The general presence of thrust ary prism due to a steady oceanic subduction from Paleozoic to Present. However, during the last 40 years, the interpretations and con- CENTRAL cepts related to the tectonic evolution of the Japanese Islands have evolved a lot and other models have been advocated. JAPAN N. Kitakami During the 1980s, alternative collisional models were proposed, taking into account the new structural figure of SW Japan, after the SWJAPAN discovery of large thrust sheets (e.g. Charvet et al., 1985; Faure et al., 1986). In this interpretation, the episodic orogenic events Hida Zone S. Kitakami are assigned to the intermittent arrival at the trench of a buoyant Tanakura feature, micro-continent or mature arc, choking the subduction CircumHida Fault Zone MTL Abukuma and inducing the piling of nappes which show a geometry similar Sangun-Maizuru to the one visible in classical collisional orogens. Similar collisional Zone anto episodes were proposed for NE Japan (Komatsu et al., 1983; Jolivet, Oga Nappe 1986; Jolivet et al., 1987; Kimura, 1997). The aim of this paper is to review, based on a brief historical TanbalItoigawa-Shizuoka Fault perspective, the validity of the main tectonic and geodynamic con- Zone cepts implied by the two opposite models, when compared with N.Chichibu bagawaNappe the recent data. I will show that the pure accretionary model is bokeUnit likely invalid and that the collision of light blocks is the only mech- Kurosegawa-Sanbosan anism able to account for the different data: tectonic, geodynamic, Zone 200km Shimanto Zone petrologic, and geochemical. A new geodynamic model, taking into account those recent data coming from the literature, is proposed Fig. 1. Sketch map showing the main divisions of Japan (modified after Faure et al., at the end. Although this reflecting is valid for the whole Japan 1987). 90 J.Charvet/JournalofAsianEarthSciences 72(2013)88-101 Hida Zone Funatsu Circum Hida Zone. granite Sangun Tsuyama Kamigori Maizur MaizuruZone Wakas Akiyosh Tanba Zone Ryoke Zone Fig.3 ObokeUnit Green Schist Nappe Superficial Nappe Kurosegawa-Sanbosan Zone Shimanto 50km Zone Fig. 2. Schematic structural map of SW Japan, emphasizing the structures of the Late Jurassic-Early Cretaceous orogen (after Faure et al., 1987). Note the tectonic windows of the Tanba zone beneath the Sangun-Maizuru units partly inherited from the Chugoku nappe system of the Late Permian-Triassic orogeny. Black spots in the Sangun-Maizuru zone represent the Triassic (T) and Jurassic (J) molasses. Fig.4 入 Wakasa N Kamigori Izumi Gr. Kumoso Ha S MTL 'SHIMANTO TANBA 50Km Tertiary- Yakuno Sangun Permian 2 volcanics ophiolite schists units Cretaceous Inferred High grade 6Upper Cretaceous Cover/Basement Sennan Gr. Low grade 10 8 Izumi Gr. volcanics basement Sanbagawa Cretaceous Tertiary 1 13. Miocene-Recent 14 Younger/older Shimanto 12= Shimanto deposits Ryoke granites ofShimanto Fig. 3. Schematic general cross-section of SW Japan (after Charvet et al., 1985). sheets and nappes in all the zones (e.g. Charvet, 1980; Faure, 1983; HP schists and Yakuno ophiolites, initially interpreted as a former Charvet et al., 1983; Faure and Charvet, 1983, 1984; Guidi et al., seamount of the subducting ocean (Caridroit et al., 1987; Faure and 1983; Caridroit et al., 1985) led to the recognition of an “Alpine- Charvet, 1987), was later proposed to be a part of the Hida-Sino- type orogen in an island-arc position" (Charvet et al., 1985). Since Korean margin (Cluzel, 1991; Chough et al., 2000). then, the structural sketch of SW Japan was changed and consid- -Do ads aos aa pe i 's) sdeu no ud e se pa 3. Recent interpretations: recognition of large thrust sheets but genic events. return to the accretionary model with ridge subduction Those major tectonic events, responsible for the Triassic, Creta- ceous, Eocene, and Miocene thrustings, were interpreted in terms q so e si seu n sia of collisional tectonics, the subduction leading to the episodic arri- Japanese geological community. The advocated allochthony is even val of buoyant micro-continents in the subduction zones (Charvet larger than previously proposed. For instance, if the North Chichibu and Faure, 1984; Faure and Charvet, 1984, 1987; Charvet et al., zone was recognized as a klippe thrust over the Sanbagawa schists 1985; Faure et al., 1987; Caridroit et al., 1987). Some variations and rooted in the Tanba zone (also named Tamba or Mino-Tanba) were suggested. For instance, regarding the Permian-Triassic belt (e.g. Faure and Charvet, 1987; Kawato et al., 1991) (Kumoso local- of Inner Japan, the limestone Oga nappe, overriding the Sangun J.Charvet/Journalof Asian Earth Sciences72(2013)88-101 91 N S Ryoke CHICHIBU NAPPE Shimanto MIKABU SANBAGAWA S. ST. NORTHERNCHICHIBU- CENTRAL SOUTHERN CHICHIBU CHICHIBU MTL ← 1·A1 Sambosan_ Kurosegawa 'BTL 1 Km Ku Sanbagawa zumi 3Shimanto 4 Low-grade 5 Central Chichibu high-grade schists schists 6Sandstone Assumed Pre-Cretaceous Chichibu Oboke-Kurosegawa 10 Green unit Sedimentary rocks nappe basement olistostrome rocks Coverof 12 Phase 1 13 Phase2 Mikabu green rocks thrust Fig. 4. Zoom on the cross-section of the “outer zones" in eastern Shikoku island (after Charvet et al., 1985). O: Oboke; Ku: Kurosegawa. Sanbagawa meta-AC Profile BTL 340Ma-WBplane 120-90 Ma WB plane (=Sanbosan AC) Osayama TL Shimanto meta-AC 450Ma-WBplane /(=N.ShimantoAC) Ng-HmTL Chichibu(= Mino-Tanba)AC 80-60Ma lore sediments Sanbosan AC WBplane ATL kiB 0km ntoAC S.Sh Oeya 10 20 Philippine Sea Plate 30 40 50 40 km Fig. 5. Simplified anatomy of SW Japan emphasizing the superimposition of accretionary complexes younging downward and the absence of basement continent-like units (after Isozaki et al., 2010). AC and meta-AC: Acretionary complexes and metamorphosed accretionary complexes; WB plane: Wadati-Benioff plane; Ng-HmTL: Nagato-Hida marginal Tectonic Line; I-KTL: Ishigaki-Kuga Tectonic Line; BTL: Butsuzo Tectonic Line; ATL: Aki Tectonic Line. the South Chichibu (Sanbosan) and the composite Kurosegawa ter- Sanbagawa schists is older than 130 Ma and the metamorphic rane, which comprises old (pre-Silurian to Siluro-Devonian) rocks age around 120-110 Ma (Aoki et al., 2007, 2008; Isozaki et al., (Yoshikura et al., 1990; Aitchison et al., 1991, 1996; Hada et al., 2010, and reference therein) or 89-88 Ma (Wallis et al., 2009). 2000) and a serpentinous melange (Isozaki and Itaya, 1991; Iso- But, if the bulk geometry of flat-lying thrust sheets of different zaki, 1996, 1997). And the Kurosegawa terrane, formerly regarded ages is widely accepted, and supported by seismic records as a part of the colliding block acting as the relative autochthon of (Kawamura et al., 2003; Ito et al., 2009), the interpretation in terms the overthrust Sanbagawa HP schists unit (e.g. Maruyama et al., of collisional orogeny is usually abandoned, at least for the post- 1984), from southern origin (Aitchison et al., 1991; Hada et al., Triassic events, to the benefit of the accretionary orogeny model. 2001), is now seen as a melange coming from the innermost zones The whole SW Japan is considered as a pile of accretionary com- (Isozaki, 1997; Maruyama, 1997; Is0zaki et al., 2010). Also, the San- plexes younging downward (Fig. 5). Indeed, since the 1980s, the use of high-resolution biostratigra- ceous North Shimanto zone in central Ki peninsula (Sasaki and phy (conodonts and radiolaria) allowed several authors to recog- Isozaki, 1992; Masago et al., 2005). And, in Oboke area, central Shi- nize ancient accretionary wedges, with a reconstructed ocean koku, the lower unit made of Koboke and Kawaguchi Formations, plate stratigraphy of individual accreted units, especially in the initially regarded as a psammitic schists unit of the HP belt, ap- Tanba and Shimanto zones (e.g. Matsuda and Isozaki, 1991; Mats- pears to be a tectonic window correlated to North Shimanto; the uoka, 1992; Is0zaki, 1996, 1997; Wakita, 2000; Wakita and Met- protolith of those formations is younger than ca. 92-82 Ma, and calfe, 2005; Isozaki et al., 2010, and reference therein). Also, the metamorphic age around 60 Ma, whereas the protolith of precise radiometric dating permitted the distinction of metamor- 92 J.Charvet/JournalofAsianEarthSciences 72(2013)88-101 phosed units, especially in Sanbagawa and Shimanto belts (e.g. Su- 4.1. Tectonic considerations zuki et al., 1990; Is0zaki and Itaya, 1991; Aoki et al., 2008, and ref- erence therein). The bulk geometry, with flat and refolded thrust contacts, is dif- As a result of these new data, in the 1990s, the plate tectonic ferent from an accretionary prism geometry. The reworking of old- interpretation of the history of the Japanese Islands was revised by er belts, in a kind of “basement nappes", implies huge deformation far beyond the back-stop. Even if it is assumed to be a simplified steadily oceanic subduction, with an episodic ridge subduction as- sketch, a cross-section like the one on Fig. 5 (after Isozaki et al., sumed to be responsible for the genesis of metamorphic-granite 2010) is misleading. The tectonic contacts between units have been belts (e.g. Is0zaki, 1996, 1997; Maruyama, 1997; Is0zaki et al., 2010). refolded and reactivated, several times for some of them; for in- The progress in understanding and modeling the exhumation of stance the Permian-Triassic stack has been thrust again at the HP metamorphic rocks, associated with precise field survey of the end of Jurassic. The actual structure does not correspond to such contacts led authors to re-interpret the Sanbagawa belt in terms of -as-jo-no, jo oe e o anq eoed ,unbas u, jo a ue a thin accretionary slice formed along the subduction channel and quence" thrusts and cannot be compared with the classical anat- emplaced by wedge extrusion (Maruyama et al., 1996; Ota et al., omy of an accretionary prism. And the main contacts cannot be 2004; Masago et al., 2005; Osozawa and Pavlis, 2007; Aoki et al., assigned to the former Wadati-Benioff planes. Actually the geome- n ) 5 ( try, involving the multiple thrusting of ancient nappe systems, attention has been given to the eclogites of central Shikoku, inter- with different metamorphic evolution, recalls the one of classical preted as an oceanic plateau in an accretionary complex (Terabay- collisional orogens. ashi et al., 2005) or derived from an oceanic island arc (Utsunomiya Nevertheless, the large allochthony advocated in some recent et al., 2011), and to their metamorphic conditions reaching locally papers can be discussed. An example is given by the Kurosegawa 2.5 to more than 3 GPa and 900 °C (Enami and Miyamoto, 2001; rocks, presented as a pre-Jurassic klippe initially thrust over the Ota et al., 2004). They have undergone two metamorphic stages Tanba-Chichibu units (Isozaki and Itaya, 1991; Isozaki et al., (Endo, 2010): a first one at around 120 Ma, the last one, eclogitic, 2010) and therefore rooted to the west of Tanba or as olistoliths at 89-88 Ma, followed by a fast exhumation, with a rate of in such a far-travelled olistostrome klippe (Masago et al., 2.5 cm/a (Wallis et al., 2009). 2005)(Fig. 6). The Kurosegawa terrane is agreed to be correlative Between 500 Ma and the Tertiary, five major orogenies occurred with South Kitakami (Fig. 1), a large massif in NE Japan with old (Isozaki et al., 2010), respectively at: 450 Ma (0eyama), 340 Ma rocks (e.g. Ehiro, 2000). It is questionable that they both represent (Renge), 240 Ma (Akiyoshi), 140-130 and 80-60 Ma. This paper is olistoliths coming from the innermost zones and reworked in a dealing essentially with the period from the Late Paleozoic to latest Mesozoic meélange; their affinity with Hida Gaien is debated. If Mesozoic-early Tertiary, during which three main tectonic events they are nappes rooted near the Hida zone (Isozaki, 1997), that occurred. would strongly argue for a collisional model. But the South Kitaka- The Late Permian-Triassic one is sealed by the Upper Triassic n e no si jss j molasse. It involved the emplacement of the Yakuno ophiolite, cum-Hida area; there is no structural evidence for that. An the ca. 240 Ma old Suo HP metamorphic rocks (part of the previous alternative and more likely interpretation is that the Kurosegawa Sangun schists now subdivided into a ca. 340 Ma Renge belt and zone, which records evidence of strike-slip faulting during the the Suo belt, dated at 253-237 Ma by Nishimura and Shibata Early Cretaceous (Kato and Saka, 2003), could be originally the x q 1 autochthonous basement of the Lower Cretaceous nappes, re- implicitly, a collisional process is advocated for the Triassic event, worked as a klippe at the end of Cretaceous, together with Chic- either by correlation with the Qinling-Dabie-Sulu suture (Oh, hibu units. This two-stage interpretation is consistent with the 2006; Isozaki et al., 2010) or due to the collision of a Proto-Japan geometry of units described in Shikoku (Kawato et al., 1991) and block with the Asian margin (de Jong et al., 2009). with the fact that the initial strike-slip faults, active during the The Early Cretaceous event, post-dated by Cretaceous basins Early Cretaceous, are cut by the younger thrust contact (Butsuzo and by the Ryoke HT metamorphism, involved the remobilization Tectonic Line) (Fig. 7) responsible for the tectonic superimposition of the Triassic belt, the piling and folding of Mino-Tanba and Chic- of those units above the Upper Cretaceous Shimanto domain (Kato hibu units. The Late Cretaceous-Paleocene event, before the depo- and Saka, 2003, 2006). sition of the Eocene Kuma Group, induced the reworking of the Early Cretaceous belt in the thrust system (final emplacement of 4.2. Geodynamic implications Chichibu-Kurosegawa klippe), the HP metamorphism of North Shimanto and Oboke schists, the exhumation of Sanbagawa blues- Any interpretative theory of the tectonic development of Japan chists and eclogites. Those two Mesozoic tectonic events are now must comply with the global geodynamic machinery acting around the world. Several geodynamic processes advocated within the except some authors advocating the collision with micro-conti- frame of the accretionary orogeny model look actually unable to nents (e.g. Otsuki, 1992). This accretionary orogen model, also build the described structures. called “Miyashiro-type orogeny”" (Maruyama, 1997) emphasizes The first and main difficulty regards the ridge subduction, con- the role of continuous oceanic and episodic ridge subduction. sidered to be responsible for the episodic tectonic crises and nappe emplacement (Maruyama, 1997; Is0zaki et al., 2010). The subduc- tion of an active oceanic ridge, according to the presently working 4. Invalidity of the accretionary orogeny: revalidation of the and well documented recent example, cannot explain the tectonic collisional model piling. Such a subduction actually induces an acceleration of the tec- tonic erosion (e.g. Behrmann et al., 1994; Bourgois et al., 1996), a However, several facts are hardly explained by an ordinary oce- collapse of the upper plate, but not an ocean-ward nappe piling. anic plate subduction. The same observation is made regarding the subduction Several lines of evidence contradict this interpretation: the ("collision") of other oceanic asperities: aseismic ridges, plateaus, geometry of structures, the comparison with the geodynamic pro- seamounts. cesses presently working around the world, and the geochemical When the subduction of an aseismic ridge participates in the evidence. shallowing of the downgoing plate and increases the coupling, it J.Charvet/Journalof AsianEarthSciences72(2013)88-101 93 N Kurosegawa N Chichibu SANBAGAWA nappe ( = Tanba) S Chichibu K1 MTL 'BTL Sa2 Sa 1 South Kitakami Kurosegawa block 1 Km Chichibu Belt N (Jurassic accretionary complex) SANBAGAWA Belt Ab-Bt zone granite-gneiss Grt zone complex (450 Ma) Ryoke Belt Chl zone 'Shimanto Belt · (Upper Cretaceous accretionary complex) · . 10 Km Charvet et al. (1985): the Kurosegawa unit is part of a previous block (South Kitakami-Kurosegawa) underthrust during the Early Cretaceous, affected by strike-slip faulting, and re-thrust during the Late Cretaceous-Paleocene event involving Sanbagawa thrusting above N Shimanto.(B) Interpretation assuming that the Kurosegawa old rocks are olistoliths in the Chichibu klippe system (after Masago et al., 2005). Sa 1, Sa 2: Lower and Upper Sanbagawa units; Ab-Bt: Albite-biotite; Olg-Bt: Oligoclase-biotite; Gr: Garnet; Chl: Chlorite. Northern Chichibu Kurosegawa Sanbagawa SouthernChichibu Mikabu ophiolite Forearc basin and/or slope deposits Strike-slip basin deposits(LowerCretaceous) Negative flower structure (Extensional duplex) strike-slip Positive flower structure xaldnp (contractional duplex) Kurosegawa rocks SHIMA.NTO ERRAN BTL :Butsuzo Tectonic Line Neogenevolcano-plutoniccomplex Fig. 7. Block-diagram showing the relationships between Shimanto, Sanbagawa, and Chichibu-Kurosegawa in eastern Ki Peninsula (after Kato and Saka, 2006). Note the faults within the Kurosegawa Terrane cut by the BTL, thrust contact with Shimanto. induces also tectonic erosion (e.g. von Huene and Ranero, 2010 and the “colliding" feature, but the process also involves accelerated reference therein); it may induce the development or the activa- tectonic erosion (von Huene and Ranero, 2010). Along Peru, tion of a fold-and-thrust belt, but in the back-arc area, mainly verg- short-term uplift during the subduction ("collision") of the ocean ing continent-ward, as documented in the Andes (e.g. Schmitz, relief was followed by subsidence, leading to formation of basins 1994; Hartley et al., 2000; McQuarrie and DeCelles, 2001; McQuar- in the middle and upper slopes, and extensional deformation (Clift rie, 2002; Muiller et al., 2002; DeCelles and Horton, 2003; Barke and et al.,2003). Lamb, 2006; Espurt et al., 2007; Gotberg et al., 2010). The crustal The sometimes advocated accretion of seamounts and/or pla- shortening is due to the bending and underthrusting of the Brazil- teaus, for instance the Carboniferous Akiyoshi-Sawadani seamount ian craton (e.g. Lyon-Caen et al., 1985), and occurs in the hot, thin chain, the Permian Maizuru seamount swarm (Maruyama et al., zone of lithospheric weakness behind the arc, expected focus of 1997; Isozaki et al., 2010), or the Late Jurassic Mikabu plateau (Iso- shortening during periods of increased compressive stress zaki et al., 1990) and Sorachi plateau in Hokkaido (Kimura, 1997) is (Cawood et al., 2009). As noticed along the Americas, the also unable to explain the development of nappes. It has been dem- subduction of ocean floor relief can uplift the land surface above onstrated in many active margins around the Pacific, and particu- 94 J. Charvet/Journal of Asian Earth Sciences 72 (2013) 88-101 larly near Japan, that the arrival of seamounts at the trench may in- ing to 2.0-2.3 GPa; the oceanic exhumation velocities for HP-LT duce some local and temporary disturbance, including some small oceanic rocks, whether sedimentary or crustal, are usually on the q r ss a nn n n si order of the mm/a (between 1 and 5 mm/a), whereas they are on trench-parallel extensional faults as they pass over the outer the order of a few cm/a in the case of continental subduction. There trench rise and that they finally subduct in the same way as the is a link between faster exhumation rates and continental subduc- surrounding oceanic lithosphere, provoking a collapse of the upper tion. The exhumation of oceanic crust is obtained only when serp- plate (e.g. Fryer and Smoot, 1985; Fryer and Hussong, 1985; entinites are present in the slab mantle. Lallemand and Chamot-Rooke, 1986; Kobayashi et al., 1987; In the case of Sanbagawa schists, they are assigned to the type B Lallemand and Le Pichon, 1987; Cadet et al., 1987; Ballance et al., protoliths (ocean-derived), as defined by Maruyama et al. (1996) 1989; Gardner et al., 2001; Fisher et al., 2004; von Huene, 2008). or, more in detail, the protolith of the eclogites was probably an Similar conclusions can be drawn in the case of oceanic pla- oceanic island arc, whereas that of the mafic schists was likely teaus, if they do not show any buoyancy contrast. Conversely, if the ocean floor MORB (Utsunomiya et al., 2011) and considered they are of great size and a bit buoyant, the disturbance at the sur- as exhumed during an oceanic subduction. face is more obvious and they may develop a conspicuous coastal However, some peculiarities lead to question this interpreta- uplift, like the Yakutat Block in Alaska (Eberhart-Phillips et al., tion, at least for the eclogitic units present in central Shikoku, cor- 2006; von Huene and Ranero, 2010). However, in the mantle, the responding to the second stage of HP metamorphism (Endo, 2010), Yakutat slab is subducting with the Pacific plate and is not moving dated at 89-88 Ma (Wallis et al., 2009). According to detailed independently as a truly distinct plate (Eberhart-Phillips et al., structural studies, eclogitic bodies form a coherent unit that over- 2006). Seismic records show that it is subducting down to lies a non-eclogitic unit with a major tectonic boundary and forms 140 km, with a slab up to 600-1000 km long (Eberhart-Phillips an eclogitic nappe (Wallis and Aoya, 2000; Aoya, 2001). et al., 2006; von Huene and Ranero, 2010). It is a composite oceanic Agard et al. (2009) already noticed that the P-T paths for San- and continental block of 15- to 20-km-thick crust (Fuis et al., 2008) bagawa rather resemble those known in well established colli- and, due to its buoyancy, its relative resistance to subduction sional cases “"despite the lack of subsequent collision". If the peak causes some deformation: the unsubducted crust becomes a fore- metamorphism reached the pressure of more than 3 GPa (Enami land fold-and-thrust belt (Fuis et al., 2008; von Huene and Ranero, - 9 s (o n 1 :o o e e n nental subduction, as well as the fast exhumation rate of 2.5 cm/a oceanic plateaus produced by LIPs, having an anomalously thick s ) s crust but with a density close or identical to the normal oceanic somehow overestimated when based on mantle-derived material, crust, their arrival at the subduction trench does not induce a ma- not implying that the crust reached that depth (Oberhansli, person. jor accretion. Regarding the biggest plateau in the world, the On- com.) and a general metamorphism peak pressure of 1.5-1.9 GPa is tong Java Plateau, covering 1900,000 km? and having a crust as more likely for the whole package of oligoclase-biotite schists, thick as 33 km, only 20% of the crust are accreted above a thrust which reached eclogite facies at a depth of 48-60 km and were ex- décollement, the lower 80% of the plateau crust are subducting humed at an average rate of only 1 mm/a if the peak metamor- (Mann and Taira, 2004). The Ontong Java Plateau-Solomon island phism occurred at 120-110 Ma (Aoki et al, 2009). Such a figure arc convergent zone is the only known example on Earth of active would be compatible with an oceanic subduction. accretion of an oceanic plateau at subduction zone. This giant is But some specific features are unusual in such a setting. The exceptional and, due to its almost entire subduction, it can be con- cluded that in general oceanic plateaus are not significant contrib- lamellae, a feature common in UHP rocks (Ota et al., 2004), and this utors to the crustal growth of arcs, and therefore, to continental quartz-eclogite is sedimentary in origin (Takasu, 1989; Ota et al., growth (Mann and Taira, 2004). 2004). Also the presence of the Higashi-Akaishi garnet-bearing In contrast, the arrival at the trench of a rather small block with ultramafic body (Mizukami and Wallis, 2005) is unique amongst a lighter crust, and therefore positive buoyancy, like a mature arc purported oceanic subduction-type metamorphic belts (Hattori or a micro-continent is able to induce a collision producing perma- et al., 2010); such lenses are known in collisional orogens. This nent compressive structures. A good and well-known example is body represents a cumulate assemblage, maybe a root of an arc provided by the Izu collision in Central Japan, even if the accretion (Hattori et al., 2010) or the base of an oceanic plateau (Terabayashi is only partial and a part of the arc is subducting (Yamamoto et al., et al.,2005) which has been subducted down to 100-120 km 2009). The thickness of continental crust exerts a major control on -o, son go g n (io e na o a ) this phenomenon, if we consider that the arcs thinner than 25 km nod s s , s () e na o subduct and the Izu arc, 30-35 km thick, induces collisional struc- the first example of a regional ultrahigh-P metamorphic belt in the tures (Yamamoto et al., 2009). Pacific-type orogens of the world, with a wide P range covering A past arc accretion-collision with an active margin likely oc- depths of subduction zone magmatism" and Hattori et al. (2010) curred at the northwestern corner of the Pacific about 55 Ma ago, assume the existence of “a rare example of oceanic-type ultra- between the Kamchatka-Koryak margin and the Olyutorsky arc high-pressure metamorphism". It seems more realistic to question (Scholl,2007). the model and admit that, at least regarding the eclogitic nappe, In conclusion, assigning the tectonic development of SW Japan the metamorphic conditions and the implied important depth to the episodic subduction of oceanic features like active ridge, sea- reached during subduction are only compatible with a continental mount chain, plateau, is not in agreement with the way plate tec- subduction. Regarding the exhumation rate, it has been recently tonics is presently working. The creation of large permanent debated (Aoki et al., 2009, 2010; Wallis et al., 2009; Wallis and ocean-ward compressive structures needs the underthrusting of Endo, 2010), interpreted as slow or rapid depending on the dating buoyant blocks. of the peak eclogitic metamorphism at 120-110 or 89-88 Ma. The latter hypothesis is supported by Lu-Hf data (Wallis et al., 2009). It 4.3. Exhumation of Sanbagawa HP metamorphic rocks is in the same range as the one of 39Ar/4°Ar dating of phengites (Dallmeyer and Takasu, 1991; Nuong et al., 2009). It is also in In their synthesis on the exhumation of oceanic blueschists and agreement with the zircon ages (mean age of 85.6 + 3 Ma) provided eclogites worldwide, Agard et al. (2009) made several statements: no exhumation is possible beyond a depth of ~70 km, correspond- dated rims as due to a retrograde stage following a 120-110 Ma J. Charvet/Journal of Asian Earth Sciences 72(2013) 88-101 95 metamorphic peak, despite the 92-104 Ma apparent age of the Ryoke and Tanba zones (e.g. Nureki and Murakami, 1979; Asami inherited magmatic zircon cores. This interpretation has been and Asami, 1982). But, as pointed out by Faure et al. (1986), if they questioned by Wallis and Endo (2010), who argued that the ca. bring evidence for a sialic basement comprising a deep continental 85 Ma age dates indeed the metamorphic peak. There is a consis- crust, and if the big amount of felsic magmatism produced during tent data set of independent methods suggesting also a HP meta- the Late Cretaceous and Paleogene sugests the existence of such a morphism around 85-90 Ma. That could imply a rapid sialic basement beneath Tanba and Ryoke (HT metamorphic equiv- exhumation (more than 2.5 cm/a) for at least some units and for alent of Tanba) zones, its presence could be due to two different them a short-lived Sanbagawa orogeny, on time scales of a few reasons. Either it is a pre-orogenic basement of Tanba-Ryoke or a million years (Wallis et al., 2009; Wallis and Endo, 2010). As they continental block underthrust during the orogeny. Owing to the are separated by tectonic contacts (Wallis and Aoya, 2000; Aoya, now generally accepted interpretation of Tanba as an accretionary 2001), all the Sanbagawa units were not necessarily metamor- complex initially formed in a setting of oceanic subduction (see phosed and exhumed at the same time and the possibility exists above), the former hypothesis can be discarded. Therefore, such that some were exhumed slowly and others quickly. That would high grade metamorphic rocks must represent the elements of a clearly add another evidence of a continental subduction setting colliding block arrived at the trench and having underthrust the for this latter fast exhumation. The final emplacement of the San- upper plate. bagawa units post-dated of course, at least slightly, the 66-61 Ma Recent geochemical studies dealing with Sr-Nd isotopic data of old blueschist metamorphism of the Northern Shimanto (Aoki granitic rocks generated in Japan from Paleozoic to Recent (Jahn, et al., 2011). 2010) bring a new and very strong support for the collisional mod- In addition, it is worth noting that the protolith of the quartz- el. The majority of the granitoids from SW Japan have high initial bearing eclogite was a sedimentary rock containing detrital zircons 87Sr/86Sr ratios, negative eNd(T) values and Proterozoic Sm-Nd with a core as old as around 1900 Ma (Okamoto et al., 2004). This model ages (Jahn, 2010). In other words the melting process pro- fact has strong implications. Either the protolith was initially ducing the magmas involved old (Proterozoic) continental crust. deposited near a relief of the oceanic bottom having a continental This is true for the Miocene Kashiwajima pluton of the Shimanto basement, before entering the subduction zone, which normally belt, which confirms the aforementioned conclusion. But this is does not fit with the advocated oceanic arc (e.g. Utsunomiya also true for the Mesozoic granitoids, namely the Upper Cretaceous et al., 2011) or a plateau (e.g. Terabayashi et al., 2005). Or, it was and Triassic ones, which post-date the early Cretaceous and the part of the accretionary prism and that implies the existence of Permian-Triassic tectonic crises respectively. Conversely, the an available Proterozoic source in Proto-Japan during the Early Cre- granitoids with a lower Sr ratio, implying a higher proportion of taceous (139-135 Ma), age of the sedimentation (Okamoto et al., mantle-derived material, were emplaced during the subduction 2004 and reference therein), precluding a composition restricted stages of our model, between two collisions. to a pile of oceanic accretionary complexes. In summary, the granitoids emplaced just after a tectonic crisis In short, some eclogite units of central Shikoku shows pressure show geochemical characteristics of post-collisional ones, quite conditions and a likely fast exhumation rate which are known so comparable with those observed in SE China and Taiwan, or in clas- far only in relation with a continental subduction, unknown in a sical collisional orogens in the European Hercynides and Caledo- setting of oceanic subduction. nides (Jahn, 2010). This fact argues in favor of the presence of a continental crust beneath the stack of nappes made of accretionary 4.4. Petrological and geochemical evidence complexes at the moment of the granitoid genesis and emplace- ment. What is the possible origin of such a continental crust? It The data coming from petrological and geochemical studies have been already used in order to show that they are supporting the result of an oceanic subduction, well documented by the ocean the prior underthrusting of a buoyant micro-continent in the case s s of the post-tectonic Middle Miocene plutonism of Shimanto (Stein possibility is the underthrusting of a continental block, responsible et al., 1994, 1996). For instance, Stein et al. (1994) stated that “a for the tectonic crisis and the nappe emplacement. simple subduction model does not explain the various magmatic affinities" and that “the proposed collision model could explain 4.5.Summary (1) the heating source, (2) the various magmatic affinities and also (3) the 0.7-0.8 GPa pressure invoking crustal thickening induced All the lines of evidence: tectonic, geodynamic, petrologic, and by the collisional event." In the view of those authors, a mantle upwelling, after collision and prior to Shikoku basin subduction, ridge subduction and support the model of the underthrusting of was likely responsible for the necessary heating source. This view a buoyant block (micro-continent or mature arc) for explaining was questioned by Shinjoe (1997), who argued that “mantle the episodic tectonic crises. Instead of the accretionary orogeny upwelling as a heat source cannot explain the strict along arc con- model most popular during the last decade, the collisional model temporaneity of the felsic magmatism and shift of the magmatism proposed during the 1980s is the only one accounting for the dif- to the Setouchi region". Actually, a slab detachment occurring after ferent data. the collision can easily explain the asthenospheric upwelling paral- lel to the belt and the along arc contemporaneity of such a near- trench magmatism. The younger Setouchi arc volcanism, more to 5.Role of tectonic erosion the north, was located at the volcanic front linked with the subse- quent Shikoku basin subduction. Another critical comment made A problem underlined by several authors is the scarcity of rem- by Shinjoe (1997) is the lack of direct evidence for the presence nants of old crust in Japan, despite the abundance of continental and collision of a micro-continent. However, indirect evidence is detrital clasts in the sedimentary rocks, the abundance at once of nse s ns outcropping Paleozoic granites not visible anymore, etc. One may 1994). add now the obvious necessity of the colliding blocks. Isozaki The presence of high grade enclaves, including granulite and not et al. (2010) pointed out that “the recent provenance analysis of in equilibrium with the Ryoke metamorphism, has also been known for a long time in the Tertiary volcanoes cross-cutting both formed in Japan and have already disappeared without evident 96 J.Charvet/JournalofAsianEarthSciences 72(2013)88-101 traces". For such reasons, the role of tectonic erosion has been block. The subsidence accompanying the incipient subduction par- advocated (Kato and Saka, 2003, 2006; Is0zaki et al., 2010). ticipated also in hiding the colliding block edge. It is well demonstrated that, during the subduction stage, con- spicuous tectonic erosion may occur, instead of accretion, the ero- 6. Origin of colliding blocks sive subduction margins representing 75% of the present active margins around the world (Scholl and von Huene, 2007, 2010). Where are the colliding blocks coming from? Without entering This erosion may result from two mechanisms. The first one is into detail, one can make some remarks. It has been suggested that the surface erosion of material of the upper plate transported to Proto-Japan was initially attached to Southeast China (Maruyama the trench and then subducted (e.g. von Huene and Cullotta, et al., 1997; Isozaki et al., 2010), as a part of a continental margin 1989; von Huene and Lallemand, 1990); it has been working out affected by rifting during the break-up of the supercontinent for the Paleozoic and Cretaceous granites (Isozaki et al., 2010). Rodinia about 750-700 Ma ago. The microcontinent was later The second one is the basal erosion. At erosional convergent mar- separated from SE China and drifted northeastward to form gins, lower plate underthrusting thins forearc crust by detaching proto-Japan. The recent isotopic geochemical data obtained in rock from the upper plate and transporting this material to the SW Japan support this evolution model as they share the same mantle. A 50 Ma period is long enough to erode the initial width characteristics as the South China Block (Jahn, 2010). They add a of the volcanic-arc and fore-arc massif of any subduction zone, new line of evidence to some previous ones coming from assuming a mean trench-volcanic-arc landward migration of paleobiogeographic considerations (e.g. Hada et al., 2001; Kido 5 km/Ma and a mean trench-volcanic-arc distance of 250 km (Lal- and Sugiyama, 2011). lemand, 1995). Evidence for basal subduction erosion of a forearc is A rifting episode is known to have occurred in SE China, creating (1) rapid (0.3-0.5 km/Ma) and substantial (3-5 km) subsidence,(2) the Nanhua rift, at about 850-800 Ma, later closed by an intracon- offshore truncation of cratonic rock, (3) retrograde (landward) tinental tectonic event around 460-450 Ma (Charvet et al., 2010 migration of the arc-magmatic front, and (4) the coastal and off- and reference therein). This rifting phase (or a younger one?), clo- shore occurrence of arc magmatic rocks (Scholl and von Huene, ser to the oceanic margin of Catahysia, may have succeeded and 2002). During the Cretaceous, the location of volcanic activity opened a new oceanic domain suggested by the around 580 Ma moved toward the Asian continent (e.g. Isozaki et al., 2010); the age of the Nomo oceanic crustal remnants. Its closure may have truncation of structures and presence of Cretaceous granite very been witnessed by the emplacement of the Ordovician Oeyama close to the Japan active trench, at a distance of about a few tens ophiolite (Igi et al., 1979; Ishiwatari, 1991; Is0zaki, 1996). of km suggest the idea of tectonic erosion which must have de- It is unclear when the other blocks, responsible for younger stroyed the hanging wall of pre-Miocene basement of Japan collisions, were detached from the main continent. Possibly, they (Yamamoto et al., 2009). were rifted later, after a collisional phase, during the subsequent As proposed by Kato and Saka (2006) for the Cretaceous history subduction stage. For instance, the “Honshu Block" was detached after collision of the South Kitakami micro-continent, tectonic ero- during the opening of the Carboniferous Yakuno oceanic domain sion likely occurred after each collisional event during the follow- and back-collided during the Late Permian (Charvet et al., 1999). ing subduction stage. This erosion was responsible for the removal Its autochthonous nature, advocated for paleobiogeographic of part of the structures and of the previously underthrusting reasons (Charvet et al., 1999), instead of a far-travelled exotic Shimomidani Maizuru Hida Renge suture Gr. Oga -Akiyoshi Honshu Yakuno Unazuki Block Early Permian x×× Unazuki sch HT met. Oga nappe Suo HP sch. Yakunoophiolite granites Late Olistostrome Permian - Middle Triassic β+ Continental clastic Limestone crust sediments Fig. 8. Schematic geodynamic model, in cross-section, for the Late Permian-Triassic orogeny, advocating a collision between the Hinterland (Hida s.l.) and a Honshu Block. J.Charvet/Journalof Asian Earth Sciences72(2013)88-101 97 block origin, is supported by the recent geochemical results (Jahn, ary between the Yeongnam massif of the southeastern Ryeongnam 2010). The timing of rifting of the other blocks is quite obscure, block and the Ogcheon belt is the sub-vertical Triassic Honam although likely Paleozoic to Mesozoic. That remains to be strike-slip fault (e.g. Cluzel, 1991; Oh, 2006). And the Yeongnam documented. massif, located to the south of the Ogcheon belt generally correlated Another still unclear point is the correlation of Hida with the with the Huanan rift of South China (e.g. Chang, 1996; Oh, 2006), is units known in China, Korea, and Russia. It has been naturally con- likely the extension of the Cathaysian part of South China (Chang, sidered as a piece of North China or Sino-Korea block because, if one 1996); only the northeastern Taebaeksan part of the Ogcheon belt looks at the pre-Sea of Japan configuration, it comes near North Kor- s. l. shows affinities with North China (Chang, 1996). In addition, ea (e.g. Kojima, 1989; Kojima et al., 2000; Yamakita and Otoh, the correlation between the Hida marginal Renge belt and the N 1999). But that implies an eastward extension of the Triassic Da- China-S China suture belt is not sure, as the HP metamorphic ages bie-Sulu-Imjingang collisional suture zone in Japan on the eastern are quite different, Carboniferous versus Triassic (e.g. Tsujimori border of Hida, within the Renge belt (e.g. Oh, 2006; Isozaki et al., et al., 2001; Oh, 2006; Isozaki et al., 2010 and reference therein). 2010) and in the Higo belt of Kyushu (Osanai et al., 2006). And, if An alternative correlation with the Asian mainland is a link with the Yeongnam massif of SE Korea is assigned initially to the North South China (e.g. Faure and Charvet, 1987; Chough et al., 2000). China Block, it is then interpreted as a major klippe thrust over A completely different category of interpretations is to consider the UHP belt (Isozaki et al., 2010). This last interpretation is difficult that the Hida-Oki Block was part of a separate Proto-Japan micro- as there is no convincing structural evidence; in Korea, the bound- continent colliding with the Asian mainland in the Late Paleozoic 150Ma Chugoku nappes P-T S Kitakami- Mino-Tanba Sanbosan Kurosegawa Torinosu T3 accretionary prism Block Tanba-N Chichibu nappes ShimantoBlock 110-120Ma T seamount Sanbosan Sanbagawa 60Ma Ku Tanba NCSC N Shimanto Izanagi-Pacific ridge Continental x+ ofcollidingblock + Accretionary Molass Fig.9. Schematic geodynamic model, in cross-section,for the 140-130 and 80-60 Ma orogenies, assigned to the two successive collisions of South Kitakami-Kurosegawa and Shimanto blocks. Note that the Izanagi-Pacific ridge arrives at the active margin after the main tectonic episode involving the exhumation of the Sanbagawa HP schists. 98 J. Charvet/Journal of Asian Earth Sciences 72(2013)88-101 (Maruyama et al., 1989) or during the Jurassic (De Jong et al., 1992) rather than at 155 Ma (Kato and Saka, 2003, 2006), reworks 2009). The latter hypothesis fits with well established data. For in- the Mino-Tanba accretionary prims as a pile of nappes and the stance, De Jong et al. (2009) include the Khanka Block in their Pro- units of the hinterland (Chugoku nappes and Hida s. l.) are thrust to-Japan; it means that the Jiamusi and Bureya blocks, which are again. At around 110-120 Ma, during a new subduction, the Kuro- similar (Zhou et al., 2010a; Wilde et al., 2010) could be also part n s sas a e s q eess of this entity. And, as a matter of fact, the rather well established water sedimentary basins and the Mino-Tanba stack is affected longitudinal correlation of the Mino-Tanba and Maizuru-Yakuno by the Ryoke HT metamorphism and granitic emplacement. The zones to the north in NE China and Russia shows all these zones North Shimanto prism is possibly beginning to grow. The subduc- bordering the eastern part of Hida-Khanka-Jiamusi-Bureya mas- tion of an oceanic realm rich in Triassic seamounts, reworked as sifs (Kojima, 1989; Faure and Natal'in, 1992; Kojima et al., 2000). olistoliths in the South Chichibu-Sanbosan sediments, is docu- The docking of this Proto-Japan with Eastern Asia would have hap- mented by arc volcanism and responsible for the Sanbagawa HP pened in the Late Triassic-Early Jurassic (De Jong et al., 2009), metamorphism. Such an oceanic realm may have included tempo- which is the time of the collision evidenced by the Heilongjiang rarily an intra-oceanic subduction, leading to an oceanic arc later HP complex running to the west of the Jiamusi massif (Wu et al., subducted and metamorphosed under eclogitic conditions (Uts- 2007; Zhou et al., 2009). On one hand, this is not incompatible with unomiya et al., 2011). This ocean is bounded to the east by a con- a common regional origin of those blocks as Jiamusi was maybe part of N Australia at around 500 Ma, together with N China, S Chi- the subduction of the fast-moving Izanagi plate (Whittaker et al., na, Tarim (Wilde et al., 2000, 2003). On the other hand, the affinity 2007a,b). The collision of this block with the active margin takes between Hida and Khanka blocks is questioned (Wilde et al., 2010) place between 80 and 60 Ma. It induces the final emplacement of and the Khanka-Jiamusi was not part of North China nor South the Sanbagawa HP schists and eclogites, the re-thrusting of the China blocks according to Zhou et al. (2010b). Kurosegawa-Chichibu domain, the basal thrust contact cutting So, the problem of the nature of Hida block is still open. To solve the former strike-slip faults, and the deformation of the Mino-Tan- it is beyond the scope of this paper and needs more work. That ba stack. The Izanagi-Pacific oceanic ridge arrives at the trench la- does not affect the interpretations presented in the next section, ter, at around 60-55 Ma, almost parallel to the margin and not in which the Hida area is shown as the hinterland, no matter it oblique as previously advocated (Whittaker et al., 2007a,b; Smith, was attached or not to the Asian mainland before the Jurassic. 2007; Muller et al., 2008; Yin, 2010). 7. Geodynamic model 8. Conclusions Taking into account the afore-mentioned remarks, a geody- The review of the evolution of ideas on the tectonic evolution of namic model can be proposed for the Late Paleozoic-Triassic and Japan during the two last decades leads to the following Cretaceous to Paleogene orogenies. It is inspired from the colli- statements. sional models previously proposed (e.g. Charvet et al., 1985, 1999; Faure et al., 1986; Faure and Charvet, 1987) with the modi- (1） The structure of SW Japan is made of a pile of sub-horizontal fications needed due to the recent discoveries mentioned before. nappes, polydeformed, with a geometry similar to the one For the Permian-Triassic Akiyoshi orogeny (Fig. 8), the scenario encountered in collisional orogens. is the following. During the Early Permian, an oceanic domain is (2) The mechanisms advocated for the tectonic building within subducting beneath another one: the Yakuno one with an anomas- the accretionary orogeny concept (Miyashiro-type orogeny, lously thick crust (Ishiwatari, 1985, 1991) and likely bordered by Nipponides, etc.) are inappropriate. Mainly, a steady oceanic an intra-oceanic arc (Suda, 2004). The Oga-Akiyoshi seamounts are located in this oceanic domain. The western end is occupied duction) of an active ridge or seamount chain is unable to by a Hida domain including the former 340 Ma high-P/T Renge belt build such structures; it induces in fact an acceleration of (e.g. Isozaki et al., 2010). This subduction leads to the approach of the tectonic erosion. the Honshu micro-continent. During the Late Permian-Early Trias- (3) Several lines of evidence suggest the episodic thrusting of a sic, the collision of the Honshu block induces the emplacement of buoyant block: micro-continent and/or mature arc. They the nappes of: HP Suo schists, younger unit of the Sangun belt (e.g. include: tectonic, geodynamic, petrologic, and geochemical Nishimura, 1998), Yakuno ophiolite, and non-metamorphosed data. Oga-Akiyoshi reefal limestone. It leads also to the deformation of (4) The three orogenic crises which took place at around 240, the Carboniferous Unazuki sediments, and the partial resedimenta- 130, and 80-60 Ma ago in SW Japan can be best explained tion of ophiolitic olistoliths in the Maizuru Group. It is followed by by a collisional model involving the successive arrival at HT metamorphism and granite emplacement in the Hida domain. the former subduction zone of the Honshu Block, the South This belt, with HP schists, is known to extend towards the south Kitakami-Kurosegawa Block, and a Shimanto Block. Those until at least Ishigaki Island (Faure and Charvet, 1987; Nishimura, features were likely previously separated from South China 1998; Nuong et al., 2008). with which they show affinities. (5) Tectonic erosion plaid likely a major role in removing mate- described (Fig. 9). At around 150 Ma, an oceanic realm is subduct- rial during the intervening subduction stages. ing beneath Chugoku domain, composed of the previous nappe (6) In the hinterland, the paleogeographic position and the affinity of the Hida block with surrounding units is still Due to that consumption, the South Kitakami-Kurosegawa conti- unclear. More work is needed to solve this question. nental block, on which the Torinosu formation is deposited, is (7) Similar tectonic features are due to similar geodynamic approaching the subduction zone, where the Mino-Tanba accre- causes all around the world. The tectonic architecture of tionary prism is built. That may explain the double, from north the Japanese Islands, resembling the one of some classical and south, supply of Precambrian clasts to the Mino sandstone collisional orogens, is not an exception and cannot be deposition area (Suzuki et al., 1991). The collision of the South Kitakami-Kurosegawa block, likely at around 140 Ma (Otsuki, as a peculiar and original model of accretionary orogeny. J. Charvet/Journal of Asian Earth Sciences 72(2013)88-101 99 Acknowledgements Charvet, J., Cadet, J.., Faure, M., Aubouin, J., 1983. Sur lmportance de la tectoniqu de nappe mesozoique au Japon central et meridional. Comptes Rendus de I'Academie des Sciences Paris 296, 1278-1286. 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J. Phys. Earth, 26, Suppl., S 301-S 307, 1978
TECTONICS OF THE RYUKYU ISLAND ARC
Koshiro KIZAKI
Geological Laboratory, University of the Ryukyus, Naha, Japan
(Received June 14, 1978; Revised August 26, 1978)
The geological and structural contrast between the north and central Ryukyus and
the south Ryukyus has been significant since the Late Mesozoic. The difference seems to
correspond to that of the nature of the Philippine Sea floor facing the Ryukyus, i.e. the Daito
Ridges and Amami plateau to the north and deeper basin to the south. The north and
central Ryukyus were a separate tectonic unit from the south Ryukyus from the Late Meso-
zoic to Middle Tertiary. Subsequently they have united to form an island arc as the
island groups shifted southeastwards with different rates in the Late Tertiary to Quater-
nary.
1. Introduction
The Ryukyu islands are a typical island arc, 1,200km long, lying between Kyushu
and Taiwan at the northwestern Pacific margin. They are composed of the Ryukyu
Trench on the Pacific side, a row of islands, a volcanic belt and the Okinawa Trough on
the continental margin. The islands are divided morphologically as well as geologically
into three groups: the north Ryukyu Osumi islands, the central Ryukyu Amami and Oki-
nawa islands, and the south Ryukyu Miyako and Yaeyama islands. They are separated
by the Tokara Channel and the Miyako Depression, which represent strike-slip fault zones.
The structural framework of the islands was studied by KONISHI (1965). His structural
zonation paralleled the zonation of the pre-Miocene basement complex of Southwest
Honshu: the Ryukyu islands were interpreted as the southwestern continuation of Outer
Zone of Southwest Japan modified into an echelon configuration by the left-lateral trans-
current dislocations of the Tokara Channel and the Miyako Depression. NAKAGAWA
(1974) correlated the island arc system of the Ryukyu islands with that of Northeast Japan.
In this paper, the author attempts to discuss the island arc system from a different point
of view. Recent investigations verified that the geological and structural discontinuity
of the north-central Ryukyus and south Ryukyu is significant. The geological and struc-
tural histories of both island groups are quite different throughout the period of Mesozoic
to Miocene. Thereafter, the island groups united into a single island arc.
2. Basement of the Islands
The basement rocks of central Ryukyu are composed of Late Paleozoic eugeosyn-
clinal sediments, including slate, chert, limestone and diabasic green rocks, whereas the
Yaeyama metamorphic rocks-phyllite, black schists and green schists including the glau-
cophane schist facies rocks-are seen only in south Ryukyu.
The Late Paleozoic eugeosynclinal sediments are correlated to the Upper Paleozoic
groups of Kyushu and Shikoku to the north and also to those of Taiwan to the south. The
structural trend of the basement group is not necessarily parallel to the general trend of the
S 301
S 302 K. KIZAKI
Fig. 1. Structural trend of the Paleozoic group in the Ryukyu Islands.
island arc but is more or less diverse (Fig. 1). The variation of the fold axes of the Paleo-
zoic group seems to have resulted from the later dislocations than the deformations of the
Mesozoic orogeny.
The Yaeyama metamorphic rocks are constructed mainly of two fold systems with
different orientations. The principal fold, which trends in NW-SE direction, has a wave-
length of several kilometers and is clearly oblique to the trend of the island arc. The
preferred orientation of minerals is parallel to the fold axis. A minor fold with EW axes
is superimposed on the principal fold. Shear zones and faults are well developed parallel
to the minor fold. Blocks of the metamorphic rocks crop out locally through the Lower
Miocene sandstone formation along the EW faults. It seems therefore that the EW fault-
ing was activated parallel to the island arc in the Middle to Late Miocene.
Paleomagnetic study of the Eocene volcanic rocks reveals a clockwise rotation of 40
of the Yaeyama islands (south Ryukyu) resulting in the present NW-SE trend of the meta-
morphic rocks (SASAJIMA, 1977). The original trend, therefore, should be in east-west
direction parallel to the present direction of the island arc.
Radiometric ages of the Yaeyama metamorphic rocks are 174my (K-Ar method),
195my (Rb-Sr method) (SHIBATA et al., 1968, 1972). The Sambagawa metamorphic
rocks representing facies similar to the Yaeyama metamorphics, have ages of 82-102my,
which are much younger than the Yaeyama rocks. But the Sangun metamorphic rocks
indicate similar ages of 159-175my. Both Sambagawa and Sangun rocks are from
Honshu, Shikoku and Kyushu. The green schist of the Tananao belt of Taiwan is re-
Tectonics of the Ryukyu Island Arc S 303
ported to have much younger ages of 82-14my (YEN, 1975). The Yaeyama metamorphic
rocks then seem to be correlated to the Sangun rocks so far as the radiometric ages are
concerned. The original sedimentary rocks are considered to be of the Late Paleozoic
age but no fossils have yet been found from the metamorphic rocks.
3. The Shimanto Belt
The terrane of the Shimanto supergroup of Late Mesozoic to Early Tertiary age con-
tinues southwards to the north-central Ryukyus from Kyushu, Shikoku and Honshu where
the supergroup signifies the geosynclinal sediments in the outer belt of Southwest Japan
on the Pacific side. The upper horizon of the supergroup in central Ryukyu is composed
of thick and coarse sandstones containing Eocene nummulites. The distribution of the
supergroup is limited to the north-central Ryukyus, whereas calcareous littoral sediments
intercalated with andesite and pyroclastic rocks of Eocene age are seen in south Ryukyu
(SHIRAO et al., 1 976). The Shimanto supergroup of the north-central Ryukyus is severely
deformed to form isoclinal folds and SE-verging thrust faults. It is also slightly metamor-
phosed to black and green phyllites and partly to green schists.
The Eocene formation of south Ryukyu shows different sedimentation facies from
that of the northern islands and no deformation and metamorphism except mere tilting and
later faulting. The area of south Ryukyu has therefore been a part of stable land mass
since the Eocene sedimentation. KONISHI (1965) stated that the Eocene basin of south
Ryukyu formed an inner stable zone against the Shimanto belt, which formed the outer
mobile zone of the islands arc. Accordingly, the islands of south Ryukyu have had to
shift more rapidly southwards so as to form the present Ryukyu island arc.
4. Miocene and Eocene Volcanism
The inner zone of the north-central Ryukyus represents a line of recent volcanic
islands, some of which are still active forming a present volcanic front. The basement of
the volcanoes is constructed usually of volcanic and pyroclastic rocks of Miocene to Plio-
cene age which are altered to show greenish color by hydrothermal solution. Thus, they
are collectively called "Green Tuff Volcanics" of Neogene age similar to those in the main
Japanese islands. Marine geological investigations demonstrated that the volcanic and
pyroclastic rocks are distributed in a 100km wide zone in the inner zone and the Okinawa
Trough of the north-central Ryukyus, but not of south Ryukyu (HONZA, 1977).
Pyroclastic rocks and andesite flows are found conformably within the Eocene littoral
sediments in south Ryukyu (SHIRAO et al., 1976). Their occurrence seems to be similar
to that of the "Green tuff" rocks, but the age is quite different. Moreover, the pyroxene
andesites of south Ryukyu show a lower alkali-lime index (58.3) than those from the Mio-
cene andesites (61.8) of the north-central Ryukyus (MATSUMOTO, 1964). The Eocene
volcanism cannot be traced to Taiwan but might be correlated with that of the Philippines.
Here again, the contrast of volcanism in age and character is remarkable between the
north-central Ryukyus and south Ryukyu (Fig. 2).
5. Distribution of the Yaeyama Group
The Yaeyama group, distributed in the Yaeyama islands of south Ryukyu, the
northern part of Taiwan and the Senkaku islands, is composed mainly of sandstone with
S 304 K. KIZAKI
Fig. 2. "Green Tuff Volcanism" in Miocene and Eocene.
intercalations of mudstone and is characterized by coal beds, cross laminae and trace fossils
signifying littoral sediments.
Palynological investigation reveals the age of the sediments to be the Lower Miocene
(TAKAHASHI and MATSUMOTO, 1964). The group is correlated with the Lower Miocene
series of Taiwan and northern Kyushu by the heavy mineral assemblage also (OBARA
and MATSUMOTO, 1964). The mineral composition is of zircon, tourmaline, garnet, and
often associated with rutile, staurolite and monazite. These minerals are probably sup-
plied from granite and gneiss, which are never found in the neighbouring islands nor Tai-
wan, but are abundant in the southeast coastal area of China. Therefore, the geological
development of south Ryukyu should be interpreted in connection with the geology of the
southeast China.
The Yaeyama group is slightly tilted structurally and faulted. This again shows
that the Yaeyama islands and their environs have been in stable conditions since Early
Miocene.
6. The Shimajiri Basin
It is not until the Latest Miocene during the deposition of the Shimajiri group, that the
north-central Ryukyus and south Ryukyu formed a common basin throughout the Ryukyu
islands.
The sediments of the group, which are composed mainly of siltstone interbedded sand-
Tectonics of the Ryukyu Island Arc S 305
Table 1. Marine geological and structural contrast of the north-central
Ryukyus to south Ryukyu around the Ryukyu islands after HONZA
(1977).
stone and tuff in the upper part, range from the Upper Miocene to the Lower Pleistocene
according to the investigation of foraminiferas (LEROY, 1964; NATORI et al., 1972; UJIIE
and OKI, 1974).
The lower portion of the group shows a deep water paleoenvironment, indicating re-
lative rise of sea level during the Uppermost Miocene to Pliocene. Then rapid regression
occurred during deposition of the upper horizon of the group on the basis of benthonic
foraminiferas as well as molluscan fossils (LEROY, 1964; NODA, 1976).
Southeastward shift of the basin occurred during the Shimajiri stage accompanied
by basement arching. The Ryukyu island arc formed in association with the uplift of the
Shimajiri group along the arc, simultaneous with the formation of the Okinawa Trough
and probably the Ryukyu Trench. This was at the end of the Lower Pleistocene.
7. Marine Geological Results
Marine geological research has been carried out around the Ryukyu islands by the
Hakurei Maru of the Geological Survey of Japan (HONZA, 1977). Honza stressed that
significant differences in the marine geology are observed between the southern part of
Ryukyu (south Ryukyu) and the northern part of Ryukyu (north-central Ryukyus). The
contrast of the geology and structure of the sea floor between the south and north-central
Ryukyus around the Ryukyu islands is shown in Table 1. It is particularly apparent in
the difference of the nature of the Philippine Sea floor facing the north-central and south
Ryukyu.
8. Tectonic Unification of the Islands
From the above considerations, it may be concluded that the geology and tectonics
of south Ryukyu differ from those of the north-central Ryukyus. The metamorphism of
the Yaeyama metamorphic rocks had completed by the end of the Jurassic time and the
S 306 K. KIZAKI
Fig. 3. Distribution of the Simanto supergroup which shows an oro-
clinal bent at the southwestern coast of Kyushu resulted from the
Kyushu western marginal shear (K.W.M.S.).
region thereafter has been in stable conditions, whereas in the north-central Ryukyus
eugeosynclinal stage has continued until the end of the Paleogene and the Early Miocene
age partly.
The ENE-WSW trend of the Shimanto supergroup changes almost sixty degrees to
the south at the western coast of Kyushu which is called "Hokusatsu Bend" (Fig. 3).
This oroclinal bend was produced by the Kyushu western marginal shear (KIZAKI, 1978)
with the left-lateral shift in the Middle Miocene age after the completion of the main
folding stage associated with thrust faulting. The clockwise rotation of south Ryukyu,
mentioned earlier, occurred in the Oligocene time.
In Taiwan, the formations including the Lower Miocene sediments are bent to the
east at the northeastern edge of the island where a right-lateral shearing of probable Middle
Miocene age has been reported (YEN, 1975).
It is therefore probable that the southeastward shift of the Ryukyu proto-islands cul-
minated in the Middle Miocene and continued thereafter, protruding relatively to the
southeast between Kyushu and Taiwan. At that time the Tokara Channel and Miyako
Depression were also activated as left-lateral faults. From these facts, it may be inferred
that the Ryukyu islands have been shifted differentially southeastwards particularly since
the Middle Miocene associated with strike-slip faults to form the modern island arc in the
Plio-Pleistocene time. Prior to the Middle Miocene, the north-central Ryukyus and south
Ryukyu had experienced different geological and structural histories.
Tectonics of the Ryukyu Island Arc S 307
I am very grateful to Dr. R.W. Murphy of Esso Exploration, Inc. and Professor S. Uyeda of Tokyo Uni-
versity for their critical review of the manuscript. Financial support for this work was provided by the Grant-
in-Aid for co-operative Research issued from the Ministry of Education (No. 234050).
REFERENCES
HONZA, E., An island arc activity in the Ryukyu arc and its difference in the northern and southern parts of
the arc, Mar. Sci., 9, 607-611, 1977.
KIZAKI, K., Kyushu western marginal shear and its significance, Earth Sci., 1979 (in press).
KONISHI, K., Geotectonic framework of the Ryukyu Islands (Nansei-shoto),J. Geol. Soc. Jpn., 71, 437-457,
1965.
LEROY, L.W., Smaller foraminifera from the Late Tertiary of Southern Okinawa, USGS. Prof. Paper, 454-F,
pp. 1-58, 1964.
MATSUMOTO, Y., Volcanic rocks of Iriomote-jima, Yaeyama Islands, Rep. Kyyshu Uniu. Exp. to the Yaeyama
Islands, Ryukyus, No. 2, 57-72, 1964.
NAKAGAWA, H., The problems on the geological structure of the Ryukyu Islands (Part 1), GDP News Lett.,
II-1-(1), Struct. Geol., No. 2, 87-92, 1974.
NATORI, H., M. FUKUDA, and M. ISHIDA, The late Cenozoic stratigraphy by plantonic foraminiferal species in
Okinawa and Miyazaki, J. Jpn. Assoc. Pet. Tech., 37, 48-53, 1972.
NODA, H., Preliminary notes on the bathyal Molluscan fossils from the Shinzato formation, Okinawa jima,
Okinawa Prefecture, southwestern Japan, Ann. Rep., Inst. Geosci., Univ. Tsukuba, No. 2, 40-41, 1976.
OBARA, J. and Y. Matsumoto, Heavy mineral assemblage of sandstones from the Yaeyama group of Iriomote-
jima, the Yaeyama Islands, Rep. Kyushu Univ. Exp. to the Yaeyama Islands, Ryukyus, No. 2, 47-56, 1964.
SASAJIMA, S., Paleomagnetism of the Eocene Series in the Ryukyu Islands and Southwest Honshu, with special
references to the Philippine Sea Plate, Mar. Sci., 9, 19-26, 1977.
SHIBATA, K., K. Konishi, and T. NOZAWA, K-Ar ages of muscovite from the crystalline schist of the northern
Ishigaki-shima, Ryukyu Islands, Bull. Geol. Surv. Jpn., 19, 529-533, 1968.
SHIBATA, K., R.K. WANLERS, H. KANO, T. YOSHIDA, T. NOZAWA, S. IGI, and K. KONISHI, Rb-Sr ages of several
so-called basement rocks in the Japanese Islands, Bull. Geol. Surv. Jpn., 23, 505-510, 1972.
SHIRAO, M., N. DOI, and H. NAKAGAWA, Geology of Ishigaki-jima in the Ryukyu Islands, Geol. Stud. Ryukyu
Isl., 1, 21-33, 1976.
TAKAHASHI, K. and Y. MATSUMOTO, Palynological study of the Yaeyama group of Iriomote-jima,the Yae-
yama Islands, Rep. Kyushu Univ. Exp. to the Yaeyama Islands, Ryukyus, No. 2, 35-46, 1964.
UJIIE, H. and K. OKI, Uppermost Miocene-Lower Pleistocene planktonic foraminifera from the Shimajiri
group of Miyako-jima, Ryukyu Islands. Mem. Natl. Sci. Mus. Tokyo, 7, 31-58, 1974,
YEN, T.P., Lithostratigraphy and geologic structure of Taiwan, Geol. Paleontol. S.E. Asia, 15, 303-323, 1975.
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36 T. Nozaka and Y . Ito Journal of Mineralogical and Petrological Sciences, Volume 106, page 36 ─50, 2011
doi:10.2465/jmps.100408
T. Nozaka, nozaka@cc.okayama-u.ac.jp Corresponding authorCleavable olivine in serpentinite mylonites from
the Oeyama ophiolite
Toshio N ozaka and Yuki I to
Department of Earth Sciences, Okayama University, Okayama 700 -8530, Japan
Olivine that has well -developed parting similar to cleavage, i.e., so -called “cleavable olivine”, occurs in perido -
tites at many localities of orogenic belts and the seafloor. Some conflicting hypotheses for the genesis of the
parting have been proposed but not yet fully proved or disproved. We present new data of structural, petrologi -
cal and mineralogical analyses of cleavable olivine and host ultramafic rocks in the Oeyama ophiolitic com -
plexes, SW Japan. The following are our key findings to understand the genesis of cleavable olivine. 1) Cleav -
able olivine is distributed in the ultramafic complexes regardless of metamorphic grade of contact aureoles. 2) Cleavable olivine from contact aureoles has variable chemical compositions by the effect of thermal metamor -
phism. 3) Cleavable olivine commonly occurs in or near serpentinite mylonites. 4) Antigorite blades commonly occur along the parting planes of olivine, and the parting planes along with antigorite blades are locally bent to the direction of foliation. 5) Poles of the parting planes of olivine tend to be distributed around a plane vertical
to the foliation of host serpentinite mylonite. From these facts we conclude that cleavable olivine was produced
during a sequence of localized plastic deformation and alteration of peridotites at temperatures around 600 °C or lower. The parting is likely to have derived from dislocation arrangement by recovery processes after plastic deformation of hydrous peridotites and have been brought into prominence during syntectonic serpentinization. The preferred orientation of the parting planes suggests that cleavable olivine is a potential indicator of regional tectonics of the upper mantle at supra
-subduction zones.
Keywords: Cleavable olivine, Parting, Deformation, Serpentinite, Peridotite, Mylonite, Oeyama ophiolite
INTRODUCTION
So-called “cleavable olivine” (Kuroda and Shimoda,
1967) is the olivine that apparently has perfect cleavage
faces parallel to (010), (100) and/or (001). Single -crystal
X-ray diffraction analyses and optical observations of
heated crystals by Aikawa (1981) have revealed that the
“cleavage faces” are actually parting planes with high dis -
location density. Because many workers have been used
the terms “cleavable olivine” and “cleavage” for the de -
scriptions of this type of olivine, we will use “cleavable olivine” as well for convenience; to avoid confusion,
however, we will not use “cleavage” for the parting in this
paper.
Since the first report from the Klamath Mountains,
United States by Hawkes (1946), cleavable olivine has been found in peridotitic rocks from many localities of orogenic belts. With its prominent appearance, cleavable
olivine has been interested by many researchers, and
some conflicting hypotheses for its genesis have been pro -posed. Kuroda and Shimoda (1967) and Kuroda (1969) suggested that the cleavable olivine was caused by high
-
pressure crystallization and resulting unusual crystal structure, but this idea have been disproved by advanced
X
-ray analyses of Aikawa (1981) and Aikawa and Toko -
nami (1987). Others ascribed the cleavable olivine to de -
formation related to tectonic movement (Velinsky and
Pinus, 1969; Aikawa, 1981; Kutty et al. 1983) or anneal -
ing after thermal metamorphism (Uda, 1984). Recently,
the occurrence of cleavable olivine has been reported
from serpentinized peridotites collected from the seafloor
(Ishii et al., 1992; Ohara and Ishii, 1998; Niida et al.,
2001; Murata et al., 2009). The lack of evidence for pro -
grade metamorphism in these peridotites and for existence
of heat source at the seafloor is inconsistent with the new -
est hypothesis that relates the formation of cleavable oliv -
ine to thermal metamorphism. In addition, observations of serpentinite mylonites from the Happo ultramafic com -
plex (Nozaka, 2005) suggested a relationship between
cleavable olivine and deformation.
The Oeyama ultramafic complex is the locality
where cleavable olivine has been most intensively studied
37 Cleavable olivine in serpentinite mylonites
and conflicting hypotheses mentioned above have been
proposed. The Wakasa ultramafic complex is another lo -
cality where cleavable olivine is frequently found as well. We, therefore, selected the two ultramafic complexes of
the Oeyama ophiolite for reinvestigation of cleavable ol -
ivine, and tested the hypotheses with a new set of struc -
tural, petrological and mineralogical data. The results of
our studies provide evidence for a close relationship of cleavable olivine to a sequence of plastic deformation and
syntectonic alteration processes.
GEOLOGICAL SETTING AND GENERAL DE -
SCRIPTIONS OF THE OEYAMA OPHIOLITE
The Oeyama ophiolite in this article refers to the ultra -
mafic complexes that are exposed from the Oeyama to Wakasa areas in close association with Renge high
-P/T
metamorphic rocks (Fig. 1a; Ishiwatari, 1989, 1990;
Nishi mura, 1998). The main components of the Oeyama
ophiolite are serpentinized peridotites, which include tec -
tonic blocks or intrusions of pyroxenite, gabbro, amphibo -
lite and jadeitite (Igi and Kuroda, 1965; Kurokawa, 1975,
1985; Kuroda et al., 1976; Uemura et al., 1979; Uda,
1984; Chihara, 1989; Yamaguchi, 1990; Tsujimori and Liou, 2004). The gabbroic rocks and amphibolites show K
-Ar radiometric ages of 400 -470 Ma (Nishimura and
Shibata, 1989; Tsujimori et al., 2000).
The Oeyama ophiolite is similar to the ultramafic
complexes exposed around the Tari -Misaka and Happo
areas (Fig. 1a) in lithology, age of amphibolite blocks and
close association with high -P/T metamorphic rocks (Arai,
1980; Nishimura, 1998; Tsujimori et al. 2000; Takeuchi,
2002). The Tari -Misaka and Happo ultramafic complexes
have characteristics suggestive of sub -arc mantle origin
(Arai and Yurimoto, 1995; Khedr and Arai, 2010), and
therefore, the Oeyama ophiolite could originate in a Pa -
leozoic supra -subduction zone (Ishiwatari, 1989; Ishiwa -
tari and Tsujimori, 2003). However, because there are some differences in radiometric ages and mineralogy of
residual peridotites between the Oeyama ophiolite and ul -
tramafic complexes around the Tari
-Misaka area, it is
possible as well that they belong to separate terranes
(Ishiwatari, 1990).
The Oeyama and Wakasa ultramafic complexes are
in fault contact with Paleozoic formations or Renge meta -
morphic rocks, and intruded by Cretaceous or Paleogene granitic rocks, and then, covered by younger sediments or
volcanics (Figs. 1b and 1c; Igi and Kuroda, 1965; Uemura
Figure 1. (a) Distribution of ultramafic complexes and the Sangun high -P/T metamorphic belt in SW Japan (Ishiwatari, 1989, 1990; Nishimura,
1998; Takeuchi, 2002). Abbreviations for ultramafic complexes: OE, Oeyama; WS, Wakasa; HP, Happo; TM, Tari -Misaka. (b) Geological
sketch map of the Wakasa area (Uemura et al, 1979). (c) Geological sketch map of the Oeyama area (Igi and Kuroda, 1965; Uda, 1984; Ku -
rokawa, 1985). Representative data of dip and strike of mylonite foliation measured in this study are also shown.
38 T. Nozaka and Y . Ito
et al., 1979; Uda, 1984; Kurokawa, 1985). The complexes
have contact aureoles with increasing metamorphic grade
toward the granitic intrusion (Fig. 2a). Details of the ther -
mal metamorphism will be described in the next section.
Lithology of original peridotite looks to be variable
between the complexes of the Oeyama ophiolite and even
within a complex. For example, estimate of modal pro -
portion of pyroxene in peridotites from the Oeyama com -
plex has been different between previous studies (Uda, 1984; Kurokawa, 1985). Incidentally, we point out that
modal compositions of original rocks must be cautiously
estimated in the case of serpentinized peridotites that con -
tain cleavable olivine, because cleavable olivine is very
similar in appearance to orthopyroxene, and serpentine
pseudomorphs after cleavable olivine may be misidenti -
fied as those after orthopyroxene (i.e., bastite) even under
the microscope. Rocks inspected in this study look to be
dominantly dunite with subordinate amounts of harzbur -
gite and lherzolite. Chromite with an equant or oval (elon -
gated by deformation) shape is common in the dunites, whereas Cr
-spinel with a vermicular shape occurs in har -
zburgites and lherzolites, showing consistency with the general tendency of a relationship between spinel mor -
phology and lithology (Arai, 1980; Matsumoto and Arai, 2001).
Serpentinite mylonites with strong foliation, which is
microscopically characterized by lepidoblastic antigorite
with olivine porphyroclasts and neoblasts, occur through -
out the Oeyama and Wakasa complexes (Fig. 2b). The
serpentinite mylonites are similar in appearance to those
from the Josephine ophiolite and the Happo ultramafic
complex (Norrell et al., 1989; Nozaka, 2005). Massive
serpentinites or peridotites without foliation occur as well,
but those in the highest
-grade metamorphic zone could be
serpentinite mylonites of which foliation disappeared by thermal metamorphism. The strikes of foliation show a general trend of NW
-SE and NE -SW in the Oeyama and
Wakasa complex, respectively, with local perturbation by faulting or folding (Figs. 1b and 1c).
ANALYTICAL PROCEDURES
Oriented samples and thin sections cut perpendicularly to foliation were prepared for structural analyses of serpenti -
nite mylonites. Samples that contain many grains of
cleavable olivine were selected for measurement of the
attitude of parting planes on a five
-axes universal stage.
We did not determine lineation, though essential for struc -
tural analysis, because of the scarcity or obscurity of lin -
eation in the inspected samples, in particular the rocks containing many grains of cleavable olivine.
The chemical compositions of the minerals were an -
alyzed using an electron probe micro
-analyzer with three
spectrometers (JEOL JXA -733) at Advanced Science Re -
search Center of Okayama University, with an accelerat -
ing voltage of 15 kV , a sample current of 10 -20 nA, and a
focused beam of 1 -2 µm in diameter. Standards used were
natural or synthetic oxides and silicates. The matrix cor -
rections employed followed the procedures of Bence and Albee (1968), using the alpha factors of Nakamura and Kushiro (1970). Representative analyses are listed in Ta -
bles 1 and 2.
Micro
-Raman spectroscopic analyses of serpentine
minerals were carried out at Department of Material Chemistry of Okayama University with a Raman spectro -
scope system (Jobin Yvon T64000) using 514.5 nm laser
excitation with 200 mW incident power, a 100 × micro -
scope objective lens, and a diffraction grating with groove
density of 1800 l/mm. Raman shift spectra of polished
thin
-sections were acquired for 30 seconds and integrated
five times. Baselines of the spectra were corrected using
GRAMSTM software.Figure 2. (a) Metamorphic zonal map showing distribution of ther -
mally metamorphic mineral assemblages. Forsterite in Zone II
includes the reticulate olivine (Fig. 3c; see text for details). Plots of the localities of mineral assemblages include those of Uda (1984). Abbreviations for minerals: Atg, antigorite; Cum, cum -
mingtonite; Di, diopside; En, enstatite; Fo, forsterite; Hbl, horn -
blende; Tlc, talc; Tr, tremolite. (b) Distribution of cleavable oliv -
ine (including its serpentine pseudomorphs) and serpentinite mylonites (including thermally metamorphosed rocks with visi -
ble foliation) in the Oeyama complex. Plots of the localities in this figure are only based on our study.
39 Cleavable olivine in serpentinite mylonites
METAMORPHIC ZONAL MAPPING
Because Uda (1984) has considered that the formation of
cleavable olivine was related to thermal metamorphism,
we examined first the effect of the metamorphism on the
ultramafic complexes and consequently confirmed the
overall descriptions of previous studies. Uda (1984) and Kurokawa (1985) divided the contact aureole of the
Oeyama complex into four or five zones. Our metamor -
phic zonal map (Fig. 2a) has no difference from those of the previous studies except for minor corrections in the
position of isograds and in index mineral assemblage of each zone.
According to the general scheme of thermally meta -
morphosed ultramafic rocks in SW Japan (Arai, 1975; Mastumoto et al., 1995; Nozaka and Shibata, 1995; Noza -
ka, 2003), each metamorphic zone is defined by the fol -
lowing mineral assemblage in the order of increasing
grade: Zone I, antigorite ± diopside; Zone II, forsterite +
antigorite ± tremolite or diopside; Zone III, forsterite + Table 1. Representative microprobe analyses of olivine
* Total iron as FeO.
OE and WS of the sample numbers indicates Oeyama and Wakasa, respectively.
40 T. Nozaka and Y . Ito
talc ± tremolite; Zone IV , forsterite + anthophyllite or
cummingtonite ± tremolite; Zone V , forsterite + enstatite
± tremolite or hornblende. Zone I is almost unaffected by thermal metamorphism. Uda (1984) has reported the oc -
currence of cummingtonite but we have not found this
amphibole, and hence, we cannot divide Zone III and
Zone IV in the zonal map (Fig. 2a). The existence of
high
-grade zones in a southern marginal portion of the
Oeyama complex could be caused by subsurface granitic intrusions (Uda, 1984). Although poor exposures in the
Wakasa complex makes detailed zonal mapping impossi -
ble, our samples have mineral assemblages corresponding to that of Zone I or II of the Oeyama complex.
Distinction between primary and metamorphic phas -
es is based on mode of occurrence and chemistry. Meta -
morphic olivine is different from primary olivine in that the former commonly shows reticulate or fine
-grained
granoblastic texture, contains tiny inclusions of magnetite
or sulfides, and has highly variable contents of forsterite component (Fo), NiO and MnO (Table 1) (e.g., Nozaka,
2003; see the next section for detail). Metamorphic ortho -
pyroxene is different from primary one in that the former shows radial
-shaped aggregation or poikiloblastic texture,
contains tiny inclusion of magnetite or sulfides and no clinopyroxene lamella, and has lower Ca and Cr contents
(Table 2) (e.g., Arai, 1975; Uda, 1984; Nozaka and Shiba -
ta, 1995). Primary minerals show evidence for deforma -
tion such as porphyroclastic texture, elongated shape, kink bands and undulatory extinction, whereas thermally
metamorphic minerals look to have formed under static
conditions.Table 2. Representative microprobe analyses of mafic minerals
* Total iron as FeO.
Nd, not determined. OE and WS of the sample numbers indicates Oeyama and Wakasa, respectively. “Ol
-parting ” means antigorite that penetrates along parting planes of cleavable olivine.
41 Cleavable olivine in serpentinite mylonites
MODE OF OCCURRENCE AND CHEMICAL
COMPOSION OF OLIVINE
We divided olivine from the Oeyama and Wakasa com -
plexes into five groups on the basis of mode of occur -
rence: 1) cleavable olivine, 2) uncleavable olivine, 3)
neoblastic olivine, 4) reticulate olivine and 5) granoblastic
olivine.
Cleavable olivine from Zone I and II commonly oc -
curs as medium -sized (1 -5 mm) prismatic, equant or por -
phyroclastic crystals in serpentinite mylonites (Figs. 3a
and b). Porphyroclasts of cleavable olivine, which are
spindle -shaped crystals elongated parallel to foliation, are
common in rocks containing neoblastic olivine (Fig. 3a), whereas cleavable olivine does not have the porphyro -
clastic shape in rocks lacking neoblastic olivine. Undula -
tory extinction is significant in some crystals but insignifi -
cant in other crystals of cleavable olivine. Although cleavable olivine occurs in some massive serpentinite
without pronounced foliation, lepidoblastic antigorite oc -curs more or less in the same samples or same outcrops. Serpentine pseudomorphs after cleavable olivine, which
can be distinguished from bastite after orthopyroxene by
parallel arrangement of antigorite blades with mesh tex -
ture or the lack of exsolution texture, are abundant in in -
tensively serpentinized rocks.
Parting planes parallel to (010) are the most domi -
nant in cleavable olivine, whereas those to (100) are sub -
ordinate and those to (001) are rare. In cases of olivine
that shows two sets of parting traces in thin section, de -
gree of cleavability is uneven: one set of parting planes exceeds the other in length, continuity and opening width. Networked fractures are common as well as parting in
cleavable olivine. It is noteworthy that at rims of some
cleavable olivine crystals, antigorite blades along with ad -
jacent olivine parting planes are bent to the orientation of
lepidoblastic antigorite in a matrix (Fig. 3b). The rims of
such olivine seem to be bent as well and fractured into
small prisms or needles (Fig. 3b). Individual crystals of
cleavable olivine from Zone I and II have uniform com -
positions except for the existence of networked, highly
magnesian part (Fig. 3c), which is reticulate olivine de -
scribed later.
Cleavable olivine from Zone III has parting planes,
though not so prominent as in Zone I and II. Kuroda and
Shimoda (1967) referred to this type of olivine as “transi -
tional stage” olivine with “indistinct cleavage”. The Zone III cleavable olivine looks an aggregate of small crystals (Fig. 3d), which appears to be somewhat similar to grano -
blastic olivine described later. However, each of the ap -
parently separated small crystals in the aggregate has elongated shape parallel to parting planes (Fig. 3d), and is
similar to adjacent small crystals in interference color and
extinction position, which, as a whole of the aggregate, are less variable than those of granoblastic olivine within
the same size of area of thin section. Small amounts of
talc and tremolite are set in the aggregate. Cleavable oliv -
ine from Zone III has minute inclusions of magnetite or sulfides and heterogeneous chemical compositions (Fig. 3e).
Uncleavable olivine, which lacks parting, occurs in
Zone I and II (Fig. 3f). This type of olivine commonly oc -
curs as medium
-sized (1 -5 mm) equant grains, showing a
slight degree of undulatory extinction. Porphyroclastic texture and coexistence with lepidoblastic antigorite are
rare. Uncleavable olivine commonly has networked frac -
tures, which are dominantly filled with lizardite in Zone I
and with antigorite in Zone II. In strongly serpentinized
rocks, serpentine (mainly lizardite) pseudomorphs after
uncleavable olivine form mesh texture.
Neoblastic olivine occurs as small (< 0.1 mm) grains
in close association with lepidoblastic antigorite in the
Table 2. (Continued)
42 T. Nozaka and Y . Ito
43 Cleavable olivine in serpentinite mylonites
matrix of serpentinite mylonites (Fig. 3a). Neoblastic ol -
ivine forms aggregates elongated parallel to foliation, or
forms tails of porphyroclastic olivine along with similarly
small -sized tremolite and chlorite. Neoblastic olivine has
homogeneous compositions and no magnetite inclusions
like cleavable and uncleavable olivine.
Reticulate olivine forms a vein or network that in -
vades or encloses crystals of cleavable, uncleavable or
neoblastic olivine, or occurs along parting planes of
cleavable olivine in Zone II (Fig. 3c). Reticulate olivine is
clearly different from the host olivine in composition.
This type of olivine commonly occurs in weakly meta -
morphosed serpentinites (Nozaka, 2003, 2005).
Granoblastic olivine occurs in Zones III -V and lo -
cally in Zone II. It commonly forms aggregates of small (0.1
-0.5 mm) grains with variable interference colors and
extinction positions, and without undulatory extinction. Its association with talc and tremolite is common in Zone III (Fig. 3g). Similarly small
-sized olivine in poikiloblas -
tic orthopyroxene in Zone V is included in this group.
Minute inclusions of magnetite with a small amount of
sulfides and heterogeneous chemical compositions are common characteristics of granoblastic olivine.
Figure 4 shows variations of NiO versus Fo contents
of olivine. Cleavable and uncleavable types of olivine
from Zone I and II have Fo contents that are variable be -
tween samples but almost uniform in each sample, regard -
less of crystal shape (i.e., porphyroclastic, prismatic or equant). Most of the cleavable and uncleavable olivine
from Zone I and II are plotted within or NiO
-poor side of
the mantle olivine array proposed by Takahashi et al.
(1987). Neoblastic olivine has compositions similar to
cleavable or uncleavable olivine from the same samples. Cleavable olivine from Zone III, reticulate olivine and
granoblastic olivine have highly variable and very differ -
ent compositions from those of cleavable, uncleavable
and neoblastic olivine from Zone I and II (Fig. 4). In each sample, the variation range of Fo and NiO content is < 0.5 mol% and 0.05
-0.1 wt%, respectively, in cleavable, un -
cleavable and neoblastic olivine from Zone I and II,
whereas it is 2 -9 mol% and 0.2 -0.4 wt% in Zone III
cleavable olivine, and reticulate and granoblastic olivine.
SERPENTINE MINERALS
In some representative samples, serpentine polymorphs
were identified with Raman scattering in the range of
200 -1100 cm−1 (Fig. 5). Among the Raman -shift peaks,
which are caused by vibration or stretching vibration of
atomic bonds, translation motion, cation substitution and an additional effect of epoxy resin used for thin
-section
preparation, the peaks around 1045 cm−1 and 370 -390
cm−1 can be used for the distinction between serpentine
polymorphs (Rinaudo et al., 2003; Groppo et al., 2006).
The results of Raman spectroscopy are consistent with the
textural characteristics of serpentine minerals, and have confirmed that mode of occurrence observed under the
Figure 3. Modes of occurrence of olivine in serpentinites and peridotites. Yellow and red arrow indicates the direction of foliation and parting
planes of olivine, respectively. (a) Photomicrograph of a serpentinite mylonite from the Wakasa complex (crossed polars). Olivine porphyro -
clasts (Ol) and neoblasts (green arrows) are set in a matrix of lepidoblastic antigorite (Atg). (b) Photomicrograph of a grain of cleavable oliv -
ine from the Wakasa complex (crossed polars). Antigorite (green arrows) formed along parting planes is bent at olivine (Ol) rim to be parallel
to the foliation defined by preferred orientation of the matrix antigorite (Atg). Networked fractures of olivine are commonly filled with lizard -
ite. (c) Back -scattered electron image of cleavable olivine from Zone II of the Oeyama complex. Irregular -shaped networks of olivine, i.e.,
reticulate olivine (green arrows) are darker than host olivine (Ol) due to its more magnesian compositions. The brightest mineral is magnetite.
(d) Photomicrograph of cleavable olivine from Zone III of the Oeyama complex (crossed polars). Talc (Tlc) and tremolite (invisible in this figure) are formed within the cleavable olivine. (e) Back
-scattered electron image of a part of Figure 3d. Olivine has heterogeneous composi -
tion and includes minute grains of magnetite (brightest particles). (f) Photomicrograph of equant uncleavable olivine (Ol) from Zone II of the Oeyama complex (crossed polars). Fractures in olivine are mainly filled with antigorite aggregates. (g) Photomicrograph of granoblastic oliv -
ine (Ol) and tremolite (Tr) from Zone III of the Oeyama complex (crossed polars). (h) Photomicrograph of a peridotite mylonite from the Happo ultramafic complex (crossed polars with gypsum plate). Porphyroclastic olivine (Ol) is penetrated by parallel blades of chlorite (Chl, bluish colors), which is partially replaced by serpentine (yellowish colors). Subgrains of the olivine porphyroclast are visible by the slight dif -
ference of interference color.Figure 4. NiO versus Fo contents of olivine. Several grains of oliv -
ine in each of 25 samples are plotted. MOA, mantle olivine array proposed by Takahashi et al. (1987).
44 T. Nozaka and Y . Ito
microscope is still useful for the identification of serpen -
tine minerals (e.g., O’Hanley, 1996).
It is antigorite that forms lepidoblastic texture in ser -
pentinite mylonites and occurs along parting planes of
cleavable olivine from Zone I and II (Figs. 5a and 5b).
These results are consistent with the description of Uda (1984) but contradict Aikawa (1981), who reported chrys -
otile along the olivine parting planes. It is evident that the blade
-shaped crystals of serpentine in our serpentinite
mylonites are antigorite, but we cannot rule out the possi -
bility that very thin films of chrysotile fill tightly opened parting spaces of olivine.
Lizardite was detected from mesh
-pseudomorphs af -
ter olivine and fillings of networked fractures of olivine in Zone I, whereas both antigorite and lizardite were detect -
ed from the mesh pseudomorphs and fracture fillings in Zone II (Figs. 5c and 5d).
STRUCTURAL RELATIONSHIP BETWEEN
CLEA V ABLE OLIVINE AND SERPENTINITE
MYLONITES
Fabric analyses of serpentinite mylonites suggest a rela -
tionship of cleavable olivine with deformation, although
the relatively coarse
-grained nature of rocks and decom -
position of olivine by serpentinization have limited the re -
liability of interpretation to some extent. Poles of domi -
nant parting planes, most of which are parallel to (010) of
olivine, tend to be concentrated in NE -SW and NW -SE
direction in the Oeyama and Wakasa complexes, respec -
tively, showing a harmony with the distribution of poles
of foliation (Fig. 6a). To avoid the effect of perturbation
by local faulting and folding, the relationship between the parting and foliation was examined in representative sam -ples. As shown by Figure 6b, the poles of parting planes
tend to be distributed around a plane perpendicular to fo -
liation.
DISCUSSION
Variation of olivine composition and effect of thermal metamorphism
Cleavable and uncleavable olivine from Zone I and II
have chemical compositions plotted within or NiO
-poor
side of the mantle olivine array (Fig. 4), suggesting that
they were originated in residual peridotites and cumulates from basaltic melts (Takahashi et al., 1987). The effect of thermal metamorphism on these types of olivine looks to
be almost lacking.
Neoblastic olivine has compositions similar to por -
phyroclasts of cleavable or uncleavable olivine in the same sample. Rocks characterized by such a bimodal dis -
tribution of grain size of olivine without significant com -
positional variations are referred to as peridotite my -
lonites, which were produced by plastic deformation
associated with dynamic recrystallization (e.g., Nozaka,
2005).
Compared with the olivine that looks to retain pri -
mary compositions, reticulate olivine and granoblastic ol -
ivine have more variable Fo and NiO contents. In agree -
ment with Uda (1984), we interpret these types of olivine as products of thermal metamorphism after serpentiniza -
tion, because they are very similar in chemical composi -
tion, mode of occurrence and spatial distribution to the
metamorphic olivine in other thermally metamorphosed
ultramafic complexes (e.g., Arai, 1975; Nozaka, 2003,
2005).Figure 5. Micro -Raman spectra of
serpentine minerals. (a) Antigorite
forming lepidoblastic texture in a
serpentinite mylonite with Zone I
mineral assemblage from the Waka -
sa complex. (b) Antigorite along
parting planes of olivine from Zone
II of the Oeyama complex. (c) Liz -
ardite forming a pseudomorph after olivine with mesh texture from
Zone II of the Oeyama complex. (d)
Antigorite forming an aggregate
that fills networked fractures of un -
cleavable olivine from Zone II of
the Oeyama complex.
45 Cleavable olivine in serpentinite mylonites
Cleavable olivine from Zone III has highly heteroge -
neous compositions similar to the reticulate and grano -
blastic olivine. It is most likely that this olivine have the
same origin as cleavable olivine in Zone I and II, and has been affected by thermal metamorphism after partial ser -
pentinization. The parting traces of cleavable olivine from Zone III are filled with thin films of serpentine, which was
formed by alteration after thermal metamorphism. Al -
though we have not found cleavable olivine that looks to
retain primary compositions in Zone III, it is common that
original chemistry of olivine is lost in similar
-grade zones
(Nozaka, 2003). This is probably a result of progress of
cation diffusion within olivine crystals during thermal
metamorphism. Kuroda and Shimoda (1967) and Uda
(1984) have found that cleavable olivine with weak part -
ing occurs in Zone II close to Zone III, but it is stressed
here that cleavable olivine actually occurs in Zone III,
where talc and tremolite occur (Fig. 3d).Figure 6. Equal area lower -hemisphere projections of poles of parting planes of olivine and foliation of serpentinite mylonites. (a) Projections
of all the measured poles of the Oeyama and Wakasa complexes. Contours represent distribution density of the poles of parting planes per 1%
area. (b) Projections of the poles in individual samples (sample number is shown at upper -left of each projection; OE and WS indicates sam -
ples from the Oeyama and Wakasa, respectively). Poles of olivine parting planes tend to be distributed around a plane (dotted great circle) that is perpendicular to the foliation surface (dashed great circle).
Conditions for mylonitization
Most of serpentinite mylonites from the Oeyama ophiolite
contain neoblastic olivine as well as lepidoblastic antig -
orite. The antigorite have X Mg(= Mg/Mg + Fe) around
0.96 or higher (Table 2). Metamorphic olivine that coex -
ists with such magnesian antigorite usually has Fo con -
tents greater than 92 mol% (Trommsdorff and Evans,
1972, 1974). Compared with the olivine coexisting with
antigorite, neoblastic olivine in the Oeyama ophiolite has
lower Fo contents (< 92 mol%), which are similar to those of primary igneous olivine (Table 1 and Fig. 4). Consis -
tently, neoblastic tremolite that coexists with the olivine neoblasts and chlorite at tails of olivine porphyroclasts is
less magnesian than granoblastic tremolite of thermally
metamorphic origin (Table 2). The neoblastic tremolite is
similar in composition to a series of retrograde tremolite,
and its relatively low Si and high Al contents suggest a
formation condition at higher temperatures than the con -
dition for thermally metamorphic tremolite as the case of
46 T. Nozaka and Y . Ito
the Happo complex (Nozaka, 2005).
From the textural and mineralogical evidence it is
considered that the serpentinite mylonites in the Oeyama
ophiolite was formed by two stages of plastic deformation
like those of the Happo ultramafic complex (Nozaka,
2005). The association of neoblastic olivine with tremolite
and chlorite suggests that the earlier deformation took place at temperatures of 700
-800 °C under hydrous con -
ditions (Nozaka, 2005). On the other hand, the formation of lepidoblastic antigorite suggests that the later deforma -
tion was associated with relatively high
-temperature (up
to 600 °C) serpentinization (Nozaka, 2005). The serpenti -
nite mylonites from the Oeyama ophiolite and the Happo
complex lack evidence for prograde regional metamor -
phism associated with deformation, and commonly show
the parallelism of structure between high -temperature (ol -
ivine + tremolite + chlorite) and low -temperature (antig -
orite + magnetite) assemblages, suggesting a sequential deformation and retrogression. Although the serpentinite mylonites have no constraint for pressure conditions, the lower
-temperature event of sequential deformation prob -
ably took place at a shallower level during a course of tectonic movement. Because more intensely mylonitized peridotites tend to be subjected to more intense overprint of high
-temperature serpentinization, the two -stage de -
formation seems to have been promoted by elevated per -
meability during exhumation of the mantle peridotites.
Formation of cleavable olivine
Hypotheses that have been proposed for the genesis of
cleavable olivine in the Oeyama ophiolite can be catego -
rized into three groups: crystallization under high pres -
sures (Kuroda and Shimoda, 1967), deformation related to tectonic movement (Aikawa, 1981) and annealing after
thermal metamorphism (Uda, 1984). Among them, the hypothesis of high
-pressure crystallization and resulting
unusual crystal structure of olivine (Kuroda and Shimoda, 1967; Kuroda, 1969) have been disproved by advanced X
-ray analyses of Aikawa (1981) and Aikawa and Toko -
nami (1987).
The main basis of the hypothesis of annealing after
thermal metamorphism is the following observations: parting planes cut metamorphic reticulate olivine; cleav -
able olivine occurs exclusively in the forsterite + antig -
orite zone (Zone II in our zonal map); and the parting
planes represent cellular walls with high dislocation den -
sity. From these observations, Uda (1984) has considered
that cleavable olivine formed in the forsterite + antigorite
zone by arrangement of dislocations during annealing af -
ter thermal metamorphism. Although Uda (1984) provid -
ed reliable data sets and made an important contribution to our understanding thermal metamorphism of serpenti -
nites, his hypothesis on cleavable olivine seems to be
based on misinterpretations as follows. 1) Metamorphic
reticulate olivine appears to be cut by parting planes but,
in fact, is cut by serpentine films that fill parting spaces
(Fig. 6 and Plate I
-2 of Uda, 1984). This microscopic re -
lationship just represents the intersection of metamorphic
olivine and parting planes, but does not suggest a tempo -
ral sequence of their formation. In addition, it is possible that the serpentine films postdated the formation of reticu -
late olivine, but this is not necessarily true if the serpen -
tine is antigorite, because temperature conditions for Zone II should allow the coexistence of olivine and antigorite.
Antigorite formed along parting planes (as shown in Fig.
3c) could predate or coexist with reticulate olivine. 2) The temperature conditions as low as 400 °C proposed by Uda
(1984) seem to be unreasonably low for the arrangement
of dislocations in olivine, because the recovery of disloca -
tion after heating are the same process as dislocation climb, which is insignificant at temperatures less than 800 °C (Toriumi and Karato, 1978; Hirth, 2002; Katayama
and Karato, 2008). 3) Even if the arrangement of disloca -
tions after thermal metamorphism was possible, we could see no reason why the metamorphic olivine, which
formed under static conditions, should have dense dislo -
cations as deformed olivine has, and why the dislocations concentrated on the same plane as a parting plane of host
olivine. 4) Uda (1984) has observed the coexistence of
cleavable and uncleavable olivine within the same sam -
ple, and has ascribed the presence and absence of parting to the difference of dislocation density between olivine crystals before thermal metamorphism; however, the
cause of this difference has not been addressed. 5) Uda
(1984) has accounted for the lack of cleavable olivine in high
-grade zones in terms of parallelism between the ol -
ivine -antigorite reaction curve and cooling paths of meta -
morphic zones (Fig. 19 of Uda, 1984). However, this ex -
planation seems to be unacceptable. Given that the main cause of the cooling of contact aureoles was thermal con -
duction, a higher
-grade zone should have a cooling rate
greater than lower -grade zones as suggested by simplified
thermal diffusion models (e.g., Jaeger, 1968). Because there is no evidence for difference in water activity be -
tween zones after thermal metamorphism, antigorite along with the parting planes of olivine should be formed in
higher
-grade zones as well as lower -grade zones if the
cause of parting was retrogression after thermal metamor -
phism.
We also present clear evidence against the hypothe -
sis that cleavable olivine was formed by processes related to thermal metamorphism. 1) As shown in Figure 2b, the
distribution of cleavable olivine looks to have no relation
47 Cleavable olivine in serpentinite mylonites
with metamorphic grade, which contradicts the observa -
tion of Uda (1984). 2) The deformation of antigorite and
cleavable olivine (Fig. 3b) suggests the earlier formation
of the parting before thermal metamorphism, by which
most of metamorphic minerals were formed under static
conditions. 3) Such a mode of occurrence and chemical
composition as shown by Figures 3d, 3e and 4 suggest the effect of thermal metamorphism on preexisting cleavable
olivine that was partially serpentinized. The obscurity of
parting of olivine in Zones III
-V is ascribed to decompo -
sition of parting space -filling antigorite by thermal meta -
morphism. In olivine crystals from high -grade zones, Uda
(1984) has found the existence of cellular structures com -
posed of walls of high dislocation density, which is simi -
lar to those found in cleavable olivine from low -grade
zones. Much more straightforward explanation for this fact than the unrealistic parallel cooling paths of meta -
morphic zones is that these dislocation structures are the remains inherited from deformation that took place before the thermal metamorphism. The dislocations in olivine could be diffused by heating during thermal metamor -
phism but this process could be less efficient, and conse -
quently, planar arrangement of dislocations was retained
more in lower
-grade zones. This is the reason why cleav -
able olivine that has retrograde serpentine films along parting traces occurs in Zone III and does not in Zone V .
The preferred orientation of the parting planes of ol -
ivine (Fig. 6) indicates a link of cleavable olivine to de -
formation. Because the parting planes of olivine com -
monly have intervening antigorite blades that are locally
bent to the direction of foliation, the parting seems to
have been generated before syntectonic serpentinization.
On the other hand, the parting does not seem to have a di -
rect genetic link with the high
-temperature deformation
by which peridotite mylonites were produced, because all
grains of cleavable olivine do not show porphyroclastic
texture. In the Happo complex, which have rocks less ser -
pentinized than those the Oeyama and Wakasa complexes,
olivine porphyroclasts in peridotite mylonites have no
clear parting. A likely explanation for all the observations is that the parting of olivine was generated in and around
localized shear zones after the formation of peridotite my -
lonites and before the formation of serpentinite mylonites.
The parting planes of olivine have high dislocation
density (Aikawa, 1981; Uda, 1984). The dominant mech -
anism of deformation of olivine under temperatures as low as 700
-800 °C, at which the mylonitization of the pe -
ridotites took place, is dislocation glide under a high dif -
ferential stress (Hirth, 2002; Katayama and Karato, 2008). The concentration of dislocations on certain planes is
caused by recovery processes after deformation (e.g.,
Toriumi and Karato, 1978). Such a plane with high dislo -cation density could be a favorable site for alteration.
However, it is evident from microscopic observations that
the parallel arrangement of parting planes of olivine does
not correspond with the distribution of subgrain boundar -
ies. An analogy to the cleavable olivine is olivine porphy -
roclasts in a peridotite mylonite from the Happo complex
(Fig. 3h). The porphyroclasts, which have subgrains but
no clear parting planes, are penetrated by parallel chlorite
blades that are slightly bent to the direction of foliation,
suggesting that a set of subparallel subgrain boundaries
provided the sites of the formation of chlorite under a dif -
ferential stress. Because the dominant parting planes of
cleavable olivine commonly have intervening antigorite
blades, it is considered that preferential alteration along a
set of parallel dislocation walls or subparallel subgrain boundaries and the penetration of growing antigorite
blades under a differential stress produced cleavable oliv -
ine with pronounced parallel parting. The dominant slip system is (010) [001] during the formation of peridotite
mylonites in the Happo complex (Nozaka, 2005). This is the probable cause of the dominance of (010) and rarity
(001) parting planes in cleavable olivine.
Figure 7 represents a schematic model of the forma -
tion of cleavable olivine through a sequence of deforma -
tion and alteration. Peridotite mylonites formed at an early
stage of deformation (Fig. 7a) in localized shear zones,
where supply of water could enhance the dislocation glide
of olivine with dominant slip system (010) [001] (Kataya -
ma and Karato, 2006, 2008). Subsequent recovery pro -
cesses yielded subgrain boundaries in porphyroclasts and
walls with high dislocation density in gently deformed crystals (Fig. 7b). These strongly or gently deformed oliv -
ine crystals are embryonic but not yet typical cleavable olivine, which has prominent parallel parting. During a subsequent stage of alteration and almost synchronous
plastic deformation, deformed olivine crystals were al -
tered along a set of parallel planes [dominantly (010) planes] with high dislocation density (dislocation walls)
and penetrated by parallel antigorite blades, and then the antigorite blades were bent parallel to foliation and grew
together with matrix lepidoblastic antigorite (Fig. 7c). Si -
multaneously, cleaved flakes of olivine probably glided and rotated with the support of lubricant antigorite. The
fabric of cleavable olivine (Fig. 6) was probably strength -
ened to some extent by the rotation of the cleaved flakes.
The rims of cleavable olivine crystals are bent as well as
the antigorite blades, but their appearance suggests a brit -
tle fracturing rather than plasticity (Fig. 3b), probably re -
flecting low temperature conditions at this stage of defor -
mation.
Syntectonic high
-temperature (around 400 -600 °C;
Nozaka, 2005) serpentinization, during which intervening
48 T. Nozaka and Y . Ito
antigorite grew and brought the parting into prominence,
was a necessary condition for the formation of cleavable
olivine. Another condition could be the original lithology of serpentinites. Kuroda and Shimoda (1967) have point -
ed out that cleavable olivine occur in dunite and does not
in harzburgite in the Oeyama complex. Most of serpenti -
nite mylonites in the Oeyama and Wakasa complexes in -
deed look to be originally dunite. Although the apparent
absence of cleavable olivine in harzburgite could just re -
flect the rarity of this type of rock in the Oeyama com -
plex, it seems to be possible as well that deformation tends to be localized in rocks richer in olivine, which is the mineral most susceptible to plastic deformation and
alteration under relatively low
-temperature conditions.
The concentration of the poles of parting planes on a
plane perpendicular to foliation (Fig. 6b) suggests the ex -
istence of a rotation axis of olivine crystals during defor -
mation. The intersections of foliation and the plane on which parting poles are concentrated (Fig. 6b) and the
general trends of foliation and parting planes (Fig. 6a)
seem to indicate the directions of shear stress. The domi -
nant direction of shearing is NE
-SW and NW -SE in the
Oeyama and Wakasa complex, respectively, and it is
somewhat oblique to the elongation axis of each complex
(Fig. 1). Such a structural fabric probably reflects regional
tectonics at the time of exhumation of the mantle rocks. In addition, asymmetric crystals of cleavable olivine (Fig.
3b) can be a useful indicator of a sense of shear.
The Oeyama ophiolite and the Happo complex are
thought to be derived from the upper mantle of a supra -
subduction zone (Ishiwatari and Tsujimori, 2003; Khedr
and Arai, 2010). The commonness of cleavable olivine at
supra -subduction zones is also suggested by recent re -
ports of its occurrence in serpentinized peridotites from the seafloor (Ishii et al., 1992; Ohara and Ishii, 1998; Ni -
ida et al., 2001; Murata et al., 2009). Cleavable olivine in these peridotites was probably formed by a similar mech -
anism. We expect that cleavable olivine in serpentinized peridotites sampled from other ophiolites and the seafloor
will be useful for further understanding of regional tec -
tonics of supra
-subduction zones.
CONCLUSIONS
Cleavable olivine was produced during a sequence of lo -
calized plastic deformation and alteration of peridotites at
temperatures around 600 °C or lower. The parting is likely
to have derived from dislocation arrangement by recovery processes after plastic deformation of hydrous peridotites and have been brought into prominence during syntecton -
ic serpentinization. The preferred orientation of the part -
ing planes suggests that cleavable olivine is a potential in -Figure 7. Schematic model of the formation of cleavable olivine during a sequence of deformation and alteration of peridotites. (a) Formation
of peridotite mylonites in localized shear zones under relatively high -temperature and hydrous conditions at an early stage of deformation.
Abbreviations: E -Ol, equant olivine; N -Ol, neoblastic olivine; P -Ol, porphyroclastic olivine. The peridotite mylonites contain chlorite and
tremolite. White arrow indicates a sense of shear. (b) Formation of subgrain boundaries or cellular walls with high dislocation density (dotted
line) by recovery processes after the early stage deformation. (c) Formation of cleavable olivine and serpentinite mylonites by syntectonic high
-temperature serpentinization. Antigorite blade (Atg) formation associated with glide and rotation of cleaved flakes of olivine widened
certain sets of parting planes. Bending of the antigorite blades along parting planes and brittle fracturing at rims of cleavable olivine took place at a late time of this stage of deformation.
49 Cleavable olivine in serpentinite mylonites
dicator of regional tectonics of the upper mantle at supra -
subduction zones.
ACKNOWLEDGMENTS
We are grateful to Professor J. Takada and Dr. M. Nakani -
shi of Okayama University for permission and assistance
for the usage of the Raman spectroscope system at their
laboratory. Our thanks also go to N. Fujiwara, S. Kondoh,
S. Nitta and Y . Yoshii for assistance in field and laboratory
works. The manuscript benefited from helpful comments
of the journal reviewers S. Arai and T. Tsujimori, and the associate editor A. Ishiwatari. This study was supported
by funding from the Wesco Scientific Promotion Founda -
tion.
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Environ Geochem Health (2010) 32:297-302 DOI 10.1007/s10653-010-9300-3 ORIGINAL PAPER Analytical survey of arsenic in geothermal waters from sites in Kyushu, Japan, and a method for removing arsenic using magnetite Kazuharu Yoshizuka · Syouhei Nishihama · Hideki Sato ① Springer Science+Business Media B.V. 2010 Abstract The objective of this study was to survey spring areas and spas are very popular as tourist the cation and anion contents of geothermal waters to destinations. In addition, there are currently (in 2009) gather fundamental information on geographical vari- 18 operational geothermal power plants, providing ations. Sixteen sites in hot spring areas on the island of 0.2% of Japan's electricity supply. Because geother- Kyushu in Japan were studied. The study focused on mal energy is a renewable resource with very low the arsenic content of the samples. Very high arsenic CO2 emissions, electricity providers are planning to concentrations (more than 0.1 mg/l) were detected in construct several geothermal power plants in the next most of the geothermal waters sampled. High contents 10 years, increasing the contribution from geothermal of boron and fluoride (more than 1.0 mg/l) were also sources to around 1% of Japan's electricity supply. detected in some samples. Arsenic removal was For two millennia, water and steam from hot springs performed on a laboratory scale using columns packed have been used for cooking by many people in Japan. with a magnetite-type adsorbent. The reduction of There is also a tradition of bathing in spas and of arsenic contamination to a concentration of less than drinking hot spring waters for medical purposes. In 0.01 mg/l could be achieved in the early stages of some geothermal waters, however, arsenic contami- adsorption (bed volume = 200). nation is more than 10 μg/l higher than the limit set by the WHO Guidelines for Drinking-Water Quality Keywords Analytical survey · Arsenic removal (2008). Geothermal water · Kyushu · Magnetite Arsenic is introduced into water by the dissolution of minerals and by concentration in groundwater (Choong et al. 2007). Geothermal water often contains relatively high contents of arsenic as a result of the leaching of arsenic from rocks, which occurs predom- Introduction inantly in the geothermal reservoir at high tempera- tures. Inorganic arsenic can occur in the environment In Japan, there are many famous hot spring areas in several forms. However, in natural waters, and, ("Onsen") located on a number of islands. These hot thus, in drinking-water, it is mostly found as trivalent arsenite or pentavalent arsenate. Drinking-water poses K. Yoshizuka () · S. Nishihama ·H. Sato 2009). Department of Chemical Engineering, The University In the last decade, much research has been done on of Kitakyushu, Hibikino 1-1, Kitakyushu 808-0135, Japan e-mail: yoshizuka @ env.kitakyu-u.ac.jp arsenic removal from contaminated groundwater, Springer 298 Environ Ge0chem Health (2010) 32:297-302 surface water, and soil using adsorption and ion- E130 E131 E132 exchange technologies (Pande et al. 1997; Choong N34 et al. 2007; Mohan and Pittman 2007; Henke 2009). Iron-based adsorbents such as zerovalent iron (Krishna et al. 2001; Zhang et al. 2004; Daus et al. 2004; Leupin (10) et al. 2005; Bang et al. 2005; Tyrovola et al. 2006; (1)-(6) Cornejo et al. 2008), iron (oxy)(hydr)oxides (Wilkie (6)-(2) and Hering 1996; Thirunavukkarasu et al. 2003a; N33 Sylvester et al. 2007; Zaspalis et al. 2007; Guan et al. 2008; Tuutijarvi et al. 2009), and their composites (12)-(14) (Joshi and Chaudhuri 1996; Campos 2002; Katsoy- iannis and Zouboulis 2002; Thirunavukkarasu et al. 2003b; Vaishya and Gupta 2004; Gu and Deng 2007; Mondal et al. 2008; Fierro et al. 2009) have frequently been used. Manganese (oxy)(hydr)oxides (Bajpai N32 and Chaudhuri 1999), aluminum (oxy)(hydr)oxides (Gregor 2001; Xu et al. 2002; Hlavay and Polyak 2005), and titanium dioxides (Fostier et al. 2008) can also be used as arsenic adsorbents. In this study, we survey the arsenic contents of 16) geothermal water samples from 16 hot spring sites in N31 Kyushu, Japan. We attempt to remove the arsenic on a laboratory scale by a column method packed with a Fig. 1 Map of geothermal waters sampling sites in Kyushu, magnetite-type adsorbent. Japan Materials and methods D-25). Samples of volume 10 cm? were filtered using a cellulose acetate membrane filter (Millipore GsWP, Analytical survey of geothermal waters 0.22 μm) to measure the cation and anion concen- in Kyushu, Japan trations. We measured the concentrations of Li+, B3+, Na+, Mg2+, Si4+, K+, and Ca2²+, and of F-, Kyushu (area: 36,732 km?; latitude: 33.45° to 30.58° Cl-, Br, NO2, NOs, and SO42-. Metal concen- N; longitude: 129.33° to 132.05° E; maximum alti- trations were measured by inductively coupled tude: 1,791 m above sea level) is an island located in plasma atomic emission spectroscopy (ICP-AES; southwestern Japan. Shimadzu model ICPS-70oo) and anion concentra- We took samples directly from the following hot spring areas in Kyushu: six sites at Beppu (1-6), three (Compact IC 761, 1 Metrohm). The total arsenic sites at Kuju (7-9), and one site at Takeda (10) in concentrations (As3+ and As5+) were measured by Oita; one site at Ureshino (11) in Saga; two sites at atomic absorption spectroscopy (AAS; Shimadzu Unzen (12-13),and one site at Obama (14）in model AA-6800 or SPCA-6210） with a hydride Nagasaki; one site at Hinaku (15) in Kumamoto; and vapor generator (HVG-1; Anezaki et al. 1999). one site at Ibusuki (16) in Kagoshima. A map of the sampling sites is shown in Fig. 1. During January to March in 2008, we collected 1 1 Removal of arsenic from geothermal waters of geothermal water directly from the fountainhead at each sampling site using a ladle. Samples were Preparation and characterization of magnetite transported to the laboratory within 8 h of sampling and kept at 25°C. Before filtration, the pH values FeCl2.4H2O (1.8 g) and FeCl3.6H2O (5.0 g) were were measured using a pH meter (Horiba model each dissolved in 20 cm? of deionized water. The Springer Environ Ge0chem Health (2010) 32:297-302 299 molar ratio of ferrous ion (Fe2+) to ferric ion (Fe3+) Results and discussion was 1/2. The aqueous solutions were mixed and aqueous sodium hydroxide solution was added to the Analytical survey of geothermal waters in mixture until the pH was 12. Magnetite was washed Kyushu, Japan with deionized water and dried at 323 K in an electric oven. The magnetite was ground for the adsorption Tables 1 and 2 show the cation and anion concen- experiments. The characteristics of the magnetite trations, the pH values, and the electrical conductiv- crystal structure were observed using a powder X-ray ities of the geothermal waters, and the temperatures diffraction meter (40 kV and 20 mA, CuKα, Rigaku of the sampling site fountainheads. The arsenic model XRD-DSC-X II). The BET surface area was concentrations in most of the samples were over ten determined from the adsorption isotherms of N2 at times the value specified by the WHO Guidelines for 77 K using an automatic specific surface area/pore Drinking-Water Quality, despite the wide range of size distribution measurement apparatus (Bel Japan pH values. The sample from Hachoubaru (site no. 7), model BELSorp-mini II). had a particularly high arsenic concentration, 3.23 mg/l, because the steam used for geothermal Adsorptionexperimentusingabatchwisemethod power generation, which is pumped up from a depth of below 1,500 m, has a high temperature (160°C) The magnetite (20 mg) and aqueous arsenic solution and pressure (0.59 MPa). (10 cm?) were mixed and shaken at 25°C for 12 h High levels of boron (WHO Guidelines value = The arsenic concentration was analyzed by AAs. The 0.5 mg/l) and fuoride (WHO Guidelines value = amount of arsenic on the adsorbent, q (mmol/g), was 1.5 mg/l) were also detected in some geothermal determined by: waters: site nos. 7, 11, 14, and 16 for boron, and site nos. 7 and 11 for fuoride. Arsenic, boron, and (Co - C)·L (1) fuoride all have to be removed from geothermal =h W waters before the waters can be used for cooking or where Co and C are the initial and equilibrium drinking. concentrations in the aqueous phase (mmol/l), respec- tively, L is the volume of aqueous solution (l), and w Removal of arsenic from geothermal waters is the weight of adsorbent (g). Because the X-ray diffraction (XRD) pattern of the Adsorption experiment using a column method magnetite was identical to that in the XRD database (Ohe et al. 2005), the magnetite could be prepared by The chromatographic operation was also conducted the co-precipitation method described. The specific using a column system. Granulated magnetite (wet surface area of the magnetite was 94.3 m2/g. volume 1.5 cm?) and glass wool were packed into the Figure 2 shows the effect of pH on the adsorption column in the form of a sandwich. The geothermal of arsenic on magnetite. The maximum adsorption of water sample from Hachoubaru (pH = 8.0) was arsenic was obtained at pH 7. This indicated that introduced into the column (fow rate: 0.2 cm?/min) magnetite can be used for arsenic removal from using a dual-plunger pump (Flom model KP-11). The geothermal waters, because the pH values of most effluent was collected with a fraction collector geothermal waters are within the range pH 3-9. (Advantec model CHF122SA), and the arsenic con- Figure 3 shows the adsorption isotherm of arsenic centration in the effluent was measured by AAS. with magnetite. From linear regression based on the The bed volume (B.V.) was calculated by: Langmuir adsorption equation (CAs/qAs Vs. CAs), arsenic adsorption was seen to progress via the B.V. = vt/V (2) adsorption of a single adsorbate onto a corresponding where v is the flow rate of the feed solution (cm?/ site on the surface of the adsorbent. The maximum min), t is the supply time of the feed solution (min), amount of arsenic adsorbed was found to be and V is the wet volume of the adsorbent (cm?). 0.95 mmol/g. Springer 300 Environ Ge0chem Health (2010) 32:297-302 Table 1 Cation concentrations, pH values, and electrical conductivity (E.C.) of geothermal waters Site no. Site name Prefecture pH E.C. (S/m) Element concentration (mg/l) As Li B Na Mg Si K Ca 1 Umi-jigoku Oita 3.8 0.404 0.21 6.0 <0.01 1,242 12.7 214 288 39.9 2 Yama-jigoku Oita 7.9 0.647 0.97 13.1 <0.01 954 4.8 145 220 24.0 3 Kamado-jigoku Oita 7.2 0.646 1.36 12.6 <0.01 1,480 2.4 203 280 42.0 4 Chinoike-jigoku Oita 2.6 0.424 <0.01 4.4 <0.01 595 27.2 98.0 128 54.6 5 Shiraike-jigoku Oita 7.1 0.151 <0.01 0.7 <0.01 198 3.1 39.0 30.9 47.2 Oniyama-jigoku Oita 6 7.9 0.641 0.88 13.4 <0.01 1,324 4.8 232 326 37.8 7 Hachoubaru Oita 8.0 0.754 3.23 9.8 33.9 1,501 0.3 328 229 16.8 8 Komatsu-jigoku Oita 5.9 0.016 0.24 <0.01 0.1 1.3 1.6 19.1 1.0 4.8 9 Sujiyu Oita 7.1 0.019 0.20 <0.01 0.4 3.0 0.9 17.0 2.0 3.3 10 Ganiyu Oita 6.3 0.138 0.33 0.6 1.5 34.0 316 75.7 83.3 3.4 11 Ureshino Saga 8.5 0.221 0.32 1.2 10.7 39.7 0.8 58.5 15.0 5.1 12 Unzen Nagasaki 2.6 0.372 0.14 <0.1 0.8 1.9 13.3 96.1 6.7 59.3 13 Unzen-kojigoku Nagasaki 3.9 0.159 0.21 <0.1 0.1 1.1 3.2 33.0 2.4 8.5 14 Obama Nagasaki 7.9 1.394 0.55 4.7 14.9 207 148 107 156 141 15 Hinagu Kumamoto 8.0 0.268 0.11 0.8 1.5 31.1 3.5 15.4 5.2 68.7 16 Ibusuki Kagoshima 6.8 0.980 0.39 1.9 6.1 131 27.0 91.9 138 258 Table 2 Anion concentrations and fountainhead temperatures of geothermal waters Site no. Site name Fountainhead temperature (°C) Element concentration (mg/l) F C1 Br NO2 NO3 SO4 1 Umi-jigoku 98 0.53 1,193 4.0 <0.2 1.6 468 Yama-jigoku 80 1.20 1,927 8.2 <0.2 2 1.9 342 3 Kamado-jigoku 90 1.32 2,220 8.6 <0.2 1.7 276 Chinoike-jigoku 78 0.61 938 3.6 <0.2 1.7 576 ? 5 Shiraike-jigoku 95 0.11 100 1.1 <0.2 3.0 187 6 Oniyama-jigoku 98 1.64 1,861 7.5 <0.2 3.8 329 7 Hachoubaru 160°C at 0.59 MPa 3.76 1,914 9.3 <0.2 <0.2 204 8 Komatsu-jigoku 98 <0.04 101 <0.1 <0.2 2.9 13.8 9 Sujiyu 60 <0.04 10.6 <0.1 <0.2 <0.2 40.1 10 Ganiyu 38 <0.04 197 <0.1 17.5 <0.2 333 11 Ureshino 53 9.23 215 <0.1 <0.2 <0.2 12.5 12 Unzen 94 <0.04 13.3 <0.1 <0.2 <0.2 1,081 13 Unzen-kojigoku 90 <0.04 8.2 <0.1 <0.2 <0.2 279 14 Obama 98 1.27 5,382 4.9 <0.2 <0.2 324 15 Hinagu 44 1.09 726 2.2 <0.2 <0.2 7.6 16 Ibusuki 72 <0.04 3,549 6.4 <0.2 <0.2 107 Figure 4 shows the breakthrough profile of arsenic was clear that, to keep the arsenic concentration from the geothermal water samples. The arsenic was below 0.01 mg/l, arsenic removal using a column slowly adsorbed, with breakthrough at B.V. = 200. It operation should stop before B.V. = 200. Springer Environ Geochem Health (2010) 32:297-302 301 0.043 3.5 0.042 3.0 0.041 2.5 0.040 w /6w] 2.0 0.039 1.5 0.037 1.0 0.036 0.5 0.035 2 4 6 8 10 12 pHeq 1000 2000 3000 4000 B.V.[-］ Fig. 2 Effect of pH on arsenic adsorption with magnetite Fig. 4 Breakthrough profiles of arsenic from geothermal waters 1.0 0.8 operation was conducted on a laboratory scale using granulated magnetite. The reduction of arsenic con- 9 centrations to less than 0.01 mg/l was obtained up to mmol/ 0.6 bed volume (B.V.) = 200. 0.4 A References 0.2 Anezaki, K., Nukatsuka, I., & Ohzeki, K. (1999). Determina- tion of arsenic(II) and total arsenic(mI, V) in water 0 2 4 6 8 samples by resin suspension graphite furnace atomic CAs [ mmol/dm3] absorption spectrometry. Analytical Sciences, 15, 829- 834. Fig. 3 Adsorption isotherm of arsenic with magnetite Bajpai, S., & Chaudhuri, M. (1999). Removal of arsenic from ground water by manganese dioxide-coated sand. Journal of Environmental Engineering, 125(8), 782-784. Conclusion Bang, S., Korfiatis, G. P., & Meng, X. (2005). Removal of arsenic from water by zero-valent iron. Journal of Haz- We surveyed 16 sites in hot spring areas in Kyushu, ardous Materials, 121(1-3), 61-67. Campos, V. (2002). The effect of carbon steel-wool in removal Japan, to measure the cation and anion concentrations of arsenic from drinking water. Environmental Geology, in the geothermal waters. Most of the geothermal 42(1),81-82. water samples contained very high levels of arsenic, Choong, T. S. Y., Chuah, T. G., Robiah, Y., Gregory Koay, F. L., & Azni, 1. (2007). Arsenic toxicity, health hazards and more than ten times the value given in the WHO removal techniques from water: an overview. Desalina- Guidelines for Drinking-Water Quality. High levels tion, 217(1-3), 139-166. of boron and fluoride were also detected in some Cornejo, L., Lienqueo, H., Arenas, M., Acarapi, J., Contreras, geothermal waters. These results show that arsenic, D., Yanez, J., et al. (2008). In field arsenic removal from n n boron, and fuoride have to be removed from tion.Environmental Pollution,156(3), 827-831. geothermal waters prior to using the waters for Daus, B., Wennrich, R., & Weiss, H. (2004). Sorption mate- cooking or drinking. rials for arsenic removal from water: A comparative The magnetite could be successfully prepared by the study. Water Research, 38(12), 2948-2954. Fierro, V., Muniz, G., Gonzalez-Sanchez, G., Ballinas, M. L., & alkali precipitation method. The adsorption of arsenic Celzard, A. (2009). Arsenic removal by iron-doped acti- on magnetite increased with pH. The highest adsorp- vated carbons prepared by ferric chloride forced hydrolysis. tion was obtained in the neutral pH region. The column Journal of Hazardous Materials, 168(1), 430-437. Springer 302 Environ Geochem Health (2010) 32:297302 Fostier, A. H., Pereira, M. S. S., Rath, S., & Guimaries, J. R. using magnetite.Journal of Chemical Engineering of terrain extends from south to north for a length of Japan, 38(8), 671676. geneous photocatalysis with TiO immobilized in PET Pande, S. P., Deshpande, L. S., Patni, P. M., & Lutade, S. L. or dredged metamorphic rocks from ocean-floor (1997). Arsenic removal studies in some ground waters of Gregor, J. (2001). Arsenic removal during conventional alu- metamorphism, and it has been widely accepted minium-based drinking-water treatment. Water Research, and Health:Part A Toxic/Hazardous Substances and 35(7), 16591664. of the primary (igneous) texture and mineralogy Gu, Z., & Deng, B. (2007). Use of iron-containing mesoporous Sylvester, P.,Westerhoff, P., Moller, T., Badruzzaman, M.,& carbon (IMC) for arsenic removal from drinking water. Boyd, O. (2007). A hybrid sorbent utilizing nanoparticles of direct knowledge on the structural and strati- hydrous iron oxide for arsenic removal from drinking water. Guan, X.-H., Wang, J.,& Chusuei, C. C. (2008). Removal of mechanism for the low-pressure ophiolite meta- arsenic from water using granular ferric hydroxide: Mac- Thirunavukkarasu, O. S., Viraraghavan, T., & Subramanian, K. roscopic and microscopic studies. Journal of Hazardous S. (2003a).Arsenic removal from drinking water using Bonatti et al., 1975; Helmstaedt, 1977; Mevel, that ocean-floor metamorphism is the probable Henke, K. (2009). Waste treatment and remediation technol- Thirunavukkarasu, O. S., Viraraghavan, T., & Subramanian, K. ogies for arsenic. In K. Henke (Ed.). Arsenic: Environ- S. (2003b). Arsenic removal from drinking water using mental chemistry，health threatsandwaste treatment (p. iron oxide-coated sand. Water, Air,and Soil Pollution, 351430). Chichester:Wiley. 142(14), 95111. Hlavay, J.,& Polyak, K.(2005). Determination of surface Tuutijarvi,T., Lu, J.,Sillanpaia,M.,& Chen, G. (2009). As(V) properties of iron hydroxide-coated alumina adsorbent adsorption on maghemite nanoparticles. Journal of Haz- Ishizuka et al.,1981,1983a, 1983b, 1984; Banno morphism, there are continuing uncertainties on revealed that the majority of the Kamuikotan Tyrovola, K., Nikolaidis, N. P., Veranis, N., Kallithrakas- Joshi, A.,& Chaudhuri, M. (1996). Removal of arsenic from In the axial zone of Hokkaido, the Kamuikotan ground water by iron oxide-coated sand. Journal of been attempted by Banno & Hatano (1963), Taza- Environmental Engineering,122(8),769771. has revealed that much of the oceanic crust has Katsoyiannis, I. A., & Zouboulis, A. I. (2002). Removal of ophiolite metamorphism. However, the fact that arsenic from contaminated water sources by sorption onto Vaishya, R. C., & Gupta, S. K. (2004). Modeling arsenic(V) iron-oxide-coated polymeric materials.Water Research, removal from water by sulfate modified iron-oxide coated 36(20), 51415155. the processes responsible for the low-pressure Krishna, M. V. B., Chandrasekaran, K., Karunasagar, D., & 645666. 1A). In the current scheme of metamorphic belts Wilkie, J. A., & Hering, J. G. (1996). Adsorption of arsenic terrains. The relatively intact suite of the ophiolite mitted to the Kyoto University, of which some of the Japanese islands, this terrain is regarded as metamorphosed ophiolites of the low-pressure ism) (Melson & van Andel, 1966; Cann & Fun- 97110. Leupin, O. X., Hug, S. J.,& Badruzzaman, A. B. M. (2005). World Health Organization (2008) Guidelines for drinking- Arsenic removal from Bangladesh tube well water with filter ki (1964), Shibakusa (1974), Herve (1975), Aguir- columns containing zerovalent iron filings and sand. Ervi- second addenda.World Health Organization,Geneva; Therefore, the Kamuikotan metamorphic rocks authors concluding that these rocks belong to the Mohan, D., & Pittman, C. U., Jr. (2007). Arsenic removal from pdf. water/wastewater using adsorbents—A critical review. Xu, Y.-H., Nakajima, T., & Ohki, A. (2002). Adsorption and cently, Imaizumi (1983) and Maekawa (1983) have removal of arsenic(V) from drinking water by aluminum- Mondal, P., Majumder, C. B.,& Mohanty, B. (2008). Effects of loaded Shirasu-zeolite.Journal of Hazardous Materials, adsorbent dose, its particle size and initial arsenic con- 92(3), 275287. centration on the removal of arsenic, iron and manganese ‘ss 1o g 0 from simulated ground water by Fe3+ impregnated acti- from contaminated water by iron oxide sorbents and porous vated carbon.Journal of Hazardous Materials,150(3) of the Kamuikotan metamorphic rocks have been 695702. Zhang, W., Singh, P., Paling, E., & Delides, S. (2004). Arsenic Ohe,K., Tagai,Y., Nakamura, S., Oshima, T.,& Baba,Y. et al., 1977). Metamorphic zonal mapping has (2005). Adsorption behavior of arsenic(III) and arsenic(V) phic facies series of the low-pressure type. The na- Springer
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yoshizuka 2010 As in geothermal waters Kyushu.txt
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J. Geomag. Geoelectr., 46,411-421, 1994
Magnetic Anomalies of Japan and Adjoining Areas
Yasukuni OKUBO1, Masahiko MAKINOI, Sigeru KASUGA2, Nobuhiro ISEZAKI3, Toshitsugu YAMAZAKI1,
Takemi ISHIHARAI, and Tadashi NAKATSUKAI
'Geological Survey of Japan, 1-1-3 Higashi, Tsukuba 305, Japan 2Hydrographic Department, Maritime Safety Agency, 5-3-1 TsukUi, Chuo-ku, Tokyo 104, Japan 3Chiba University, 1-33 Toyoi-cho, Inage-ku, Chiba 263, Japan
(Received November 6, 1992; Revised May 17, 1993; Accepted June 21, 1993)
The magnetic anomaly map of Japan and adjoining areas was produced by compiling many data sets
collected by air-borne and ship-borne surveys. In order to tie between adjoining areas, IGRF removal,
upward continuation, and a linear shift were applied to each data set.
Linear features are detected from the map as a special reference to suggest possible crustal
structures.
Two significant features in the linear map are found. The one is the existence of lineations crossing
the trench. The lineations are caused by oceanic magnetic layers that lie beneath the overriding plate. The
distance from trench to the end of lineation is variant. The variation of distance depends on the amplitude
of anomaly, the structure of the topmost part of lithosphere, the dip angle of subducting plate, and the
Curie point depth.
The second is the occurrence of two types of lineations found in the back-arc spreading regions
described as, (1) lineation& running parallel with the arc trend on a rifting or spreading stage, and (2)
lineations across the arc trend in the young back-arc rift region. They are related to igneous rocks
produced by volcanic activities of back-arc rifting and spreading. Consequently, the two features express
igneous rocks either (1) parallel with, or (2) across the arc trend. They are probably associated with two
different spreading stages of the back-arc basin.
1. Introduction
One of the most significant results of the past magnetic surveys in geophysical research is
identification of magnetic lineations recorded on the oceanic crust. The result yields the Vine-Matthews
model (Vine and Matthews, 1963) of the magnetic structure of the igneous oceanic crust and confirms the
ocean-floor spreading hypothesis. Since then, a great number of magnetic surveys have been carried out
to the far reaches of the globe and have propelled to create crustal models.
Spreading rate and age of oceanic plate have been estimated chiefly from the alignment and lateral
dimension of magnetic lineation in an oceanic plate. Transform faults are inferred from the discontinuity
of magnetic lineations. Faults even in the continental crust can be detected from magnetic lineations (e.g.
Brabb and Hanna, 1981). Evidences of oceanic plate subduction can be recognized from the vanishing of
linear anomalies across the trench. Thus magnetic linear features reflect crustal structures and movements.
During recent years, efforts to compile the data in and around the Japanese islands collected by
several organizations, have been made (e.g. Makino et al., 1992a, b; Geological Survey of Japan and
Committee for Co-ordination of Joint Prospecting for Mineral Resources in Asian Offshore Areas, 1993).
The maps demonstrate the continuity of the magnetic lineations regarding ocean-floor and marginal sea
spreading, and show clearly the relation between the magnetic anomalies and the static and dynamic
crustal models. Particular examples are lineations in the Japan Sea (Nakasa and Kinoshita, 1994), and the
Shikoku basin (Okino et al., 1994), where correlations with the reversal time scale have been difficult
mainly because of insufficient data coverage.
411
412 Y. OKUBO et al.
=i t yr ' /, l~ t/AfM/1M 1 ~ - .` a a
~ r A ~y
rt ~
1 :~ t
r .wsn j
i
S h {
1
Fig. 1. Magnetic anomaly map of the Japanese Islands and adjoining areas (Makino et al., 1992a). Total intensity anomalies in
nTesla at a constant level of 3200 m above sea level. Data are mainly from the Geological Survey of Japan, the Hydrographic
Department of Maritime Safety Agency, the New Energy and Industrial Technology Development Organization, Chiba
University, and the University of Tokyo.
At first, authors introduce the way to make the map produced by Makino et al. (1992a) shown in Fig.
1. The most remarkable feature is that the map is covered by a great number of linear trending anomalies.
Most of them are magnetic lineations of oceanic crust. Some of them lie in the continental crust. We choose
significant linear features from the map and interpret them subsequently.
2. Summary of Data Acquisition and Processing
The subject reported in this chapter concerns the processing of magnetic data used to compile the
magnetic anomaly map of the Japanese Islands and adjoining areas (Fig. 1).
Total intensity magnetic anomaly data used in this compilation were collected by the aircraft chiefly
on land area and the ship on offshore area. The air-borne magnetic data were provided by the Geological
Magnetic Anomalies of Japan and Adjoining Areas 413
Survey of Japan (GSJ) and the New Energy and Industrial Technology Development Organization
(NEDO). The marine data were collected with much effort of scientists of the Japan Hydrographic
Department of Maritime Safety Agency, GSJ, Chiba University, the University of Tokyo, and other
institutes. A large number of data sets were offered by the Japan Oceanographic Data Center which is in
charge to bank the data provided by other institutes. All tracks of ship-borne survey are displayed in Fig.
2. The interval of track of the air-borne magnetic survey ranges 1 km to 4 km. Detail specification of the
survey appears in Makino et al. (1992b).
We begin with compilation ofthe air-borne magnetic data. The air-borne magnetic data sets provided
by GSJ are mainly from offshore areas with the exception of Hokkaido, while the NEDO data covers the
remaining land area. Most of the area covered by GSJ was flown at a constant barometric altitude of 457
m (1500 feet). The data collected by NEDO was generally recorded at 1372 m (4500 feet) constant
barometric altitude, however, there were many deviations from this altitude over the higher mountainous
areas of Japan. All flight line data were converted to grid. GSJ data were upward continued and merged
into the NEDO data to create six data sets of Hokkaido, Tohoku, Kanto and Tokai, Chubu, Chugoku and
Shikoku, and Kyushu (Fig. 3, Okubo et al., 1985). Since the altitudes were different, all data were
differentially continued and tied to common and constant datum to produce a continuous air-borne map.
A common datum of 3200 m was used since this was the highest altitude used in the previous continuation
processes. The completed air-borne map of the Japanese Islands was published in a scale of 1:2,000,000
(Makino et al., 1992b).
Large ship-borne data collection and compilation around Japan have been conducted by the
Hydrographic Department since 1967. Geomagnetic total intensity anomaly data of the continental shelf
around Japan were merged into a continuous map on the sea level (Oshima, 1987a). The map was
i 0
O --
-i -
1
Fig. 2. Track of ship-borne survey (Makino et al., 1992a).
414 Y. OKUBo et al.
Hokkaido
Chubu ohoku
~' Kanto and Tokai
Chugoku and Shikoku
`K
yushu
Fig. 3. Index map of the six areas of air-borne magnetic data compilation. The entire area was divided into six areas on the way
to making a continuous air-borne magnetic map of the Japanese Islands.
Onshore area Offshore area
Air-borne GSJ data Air-borne NEDO data Ship-borne HDJ,
(IGRF residual) (IGRF residual) GSJ, other data (IGRF residual)
Upward continue
and merge
6 data sets Upward continue
(Hokkaido, Tohoku, Kanto & Tokai, to the highest
Chubu, Chugoku & Shikoku, Kyushu) altitude
Upward continue Remove
to the highest linear shift
altitude and tie
Air-borne map of
the Japanese Islands
(1:2,000,000)
Merge at area boundaries
Magnetic map of the Japanese Islands and adjoining areas
(1:5,000,000)
Fig. 4. Data processing flow to make the magnetic map. GSJ, NEDO and HDJ denote Geological Survey of Japan, New Energy
and Industrial Technology Development Organization, and Hydrographic Department of Maritime Safety Agency, respec-
tively.
published in a scale of 1:3,000,000 (Hydrographic Department, 1983). The compiled magnetic map of
southern sea of Japan was made by Kasuga et al. (1992) in a scale of 1:6,000,000.
The air-borne and ship-borne data were taken from the International Geomagnetic Reference Field
(IGRF) residual magnetic data over each area. IGRF used was taken from the model for data collection
date of each survey. The ship-borne data were upward continued to the altitude of air-borne map (3200
Magnetic Anomalies of Japan and Adjoining Areas 415
m). Since different processes were used to compile, it is necessary to reprocess the ship-borne data sets.
A linear shift was applied to each data set to improve the tie between adjoining areas. Finally, a weighted
averaging procedure was used to merge data at area boundaries. Figure 4 is the flow of entire processes.
The data were transformed to the coordinate system using Lambert azimuthal equal-area projection
and gridded at 2.5 km interval using minimum curvature method.
Figure 1 is the map of magnetic anomalies of Japan and adjoining areas published in a scale of
1:5,000,000 (Makino et al., 1992a). The map was plotted in color with coast lines, main rivers, and
bathymetry.
3. Linear Features of Magnetic Anomalies of Japan and Adjoining Areas
The magnetic map of the Japan and adjoining areas displays a great number of linear trending
anomalies. Sources of the anomalies are mainly magnetic layers of oceanic plates, ridges that produce
oceanic plates, and igneous rocks generated by island arc volcanic activities in overriding plates. Since
they have often been deformed by tectonic movements, evolution of structures, reversely, can be
speculated from the configuration of linear anomalies. We focus on choosing linear magnetic anomalies
and interpreting their origins.
Figure 5 shows magnetic lineations drawn from the map of Fig. 1. Since, in the northern hemisphere
in general, normal magnetized sources produce dipole anomalies of a high over the south and a low over
the north and all lineations were drawn through the peak of high anomalies that extend longer than about
50 km, several lineations may shift several kilometers south off the sources.
There is a possibility that a precise study with the help of other data bears another lineation map,
because the lineations of Fig. 5 are drawn only from the compiled map of Fig. 1. In fact, lineations in the
Shikoku basin proposed by Okino et al. (1994) and in the Japan Sea proposed by Nakasa and Kinoshita
(1994) differ from Fig. 5. The lineation map of Fig. 5 resulted to be a reference of linear anomaly that can
be made from the map of 1:5,000,000.
The lineations are interpreted on the two areas: (1) land area of the Japanese Islands, and (2) offshore
area.
3.1 Land area
On the land area, four highly linear positive magnetic belts are discernible (anomalies marked A, B,
C, D on Fig. 5). Honkura et al. (1991) attempted to combine the two magnetic belts (A and D in Fig. 5)
and NE-SW trending belt (E in Fig. 5) into one linear belt. The four belts were interpreted in the
explanatory text of the magnetic map of the Japanese Islands (Makino et al., 1992b).
Magnetic belt A
Possible sources of the magnetic belt (A in Fig. 5), which has a north-south trend across Hokkaido
and the east coast of Tohoku district, are (1) the volcanics of the Rebun-Kabato Belt (Ogawa and Suyama,
1975), (2) the Kitakami granitic rocks (Finn, 1994), or (3) intrusives of ultra-basic rocks (Segawa and
Oshima, 1975).
Magnetic belt B
The magnetic belt B in Fig. 5 is largely associated with the north-south trending Kamuikotan Belt
of the Sorachi-Yezo groups. The Kamuikotan Belt consists of a large amount of serpentinite, basic
volcanics, and sediments. Coincidence of the magnetic belt with the ultra mafic rocks of the Kamuikotan
Belt suggests that these mafic rocks are the main source of the magnetic anomaly.
Magnetic belt C
Though the magnetic belt C in Fig. 5 is not fully shown on the map of Fig. 1, it is commonly known
that this belt extends continuously along the fore-arc basin of the Kuril arc (Solov'yew and Gainanov,
1963; Uyeda et al., 1967; Segawa and Oshima, 1975; Grapes, 1986). Magnetic analysis and result of
seismic survey suggest that the source of the magnetic belt is basic to ultra basic rocks related to the
volcanics beneath the thick Neogene to Cretaceous sedimentary basin.
416 Y. OKUBO et al.
45°N+ + +
Hokkai O
B ,r
C
1
A
+ +
40°N+
Japan Sea
0 100 200 300 400 500 km
7 ../
` 00, 01-010 000
+ Q 35°N+ SoU + Honshu + O Q
E S ikok
0 Kyushu / 000, d / an,lro `Hachiio rn c
10
30°N+
shinosh~
.00
00,
__-2 5~dito o C( 5ZE 130°E `/~ i ....~ 145°E 135°E 14
Fig. 5. Magnetic lineations. Lineations are drawn from the magnetic map of the Japanese Islands and adjoining areas of
1:5,000,000 that is identical with Fig. 1.
Magnetic belt D
In the west of Honshu, there is an east-west trending lineation (D in Fig. 5). The lineation forms a
belt and the belt is probably caused by the magnetite-series granitic rocks, Miocene volcanics, and
Quaternary volcanics exposed along its length. The belt is terminated by the Sangun-Renge Belt along its
northern edge. Therefore, the extent of the lineation indicates that the intense volcanism exposed in places
along its length, originally formed an elongated belt.
3.2 Offshore area
Major sources of lineations of the offshore area are magnetic layers of oceanic plate which were
Magnetic Anomalies of Japan and Adjoining Areas 417
produced at ridges, and ridges themselves. An example for the latter case is lineation F (see Fig. 5) at the
west margin of the Shikoku basin. This is caused by Kyushu-Palau ridge.
Continuous magnetic lineations over the Pacific plate which had been noted were delineated by
Uyeda et al. (1967). To construct evolution model of the oceanic plate, the ages of their sources were
deduced by many authors (e.g. Larson and Chase, 1972; Tamaki and Larson, 1988; Nakanishi et al., 1989,
1992).
In the Shikoku basin, linear anomalies were identified by Watts et al. (1977). From the lineations,
they determined aspects of the tectonic evolution of the basin. Okino et al. (1994) re-examined the
evolution from the data recently obtained.
The features of the lineations are related to their origins. They also have often been distorted by
subsequent deformation. From these aspects, three features are significant in the lineations of the offshore
area: (1) Lineations deformed by subduction, (2) lineations parallel to the arc trend in the backarc basin,
and (3) lineations perpendicular to the arc trend in the backarc basin.
Lineations deformed by subduction
Several lineations shown in Fig. 5 cross the trench. This demonstrates that magnetic subducting
oceanic layers lie beneath non- or weak magnetized overriding crusts. Along the Japan trench, the western
limit of the lineations extending from the Pacific plate is situated at 100 km away from the trench. The
other ends are at about 80 km away from the Izu-Ogasawara trench. In the Philippine Sea plate, the
northwestern limit is located at about 80 km away from the Nankai trough and at about 50 km away from
the Ryukyu trench. The variation of distance depends on the amplitude of anomaly, the dip angle of
subducting plate (Oshima, 1987b), structure of the topmost part of lithosphere (Kinoshita et al., 1986;
Kinoshita and Matsuda, 1989) and the Curie point depth where magnetic sources loose their intense
magnetization (Okubo et al., 1991 a).
Lineations parallel to the arc trend in the back-arc basin
Back-arc rifting and spreading are caused by a volcanic activity in a back-arc basin. Young rift system
of the Izu-Ogasawara (Bonin) arc lies west of the volcanic front (Karig and Moore, 1975; Tamaki and
Miyazaki, 1984). In the Okinawa trough, several research works suggest evidences of back-arc rift system
(Yasui et al., 1970; Ishihara and Murakami, 1976; Herman et al., 1978; Lu et al., 1981). The Japan Sea
is an inactive back-arc basin which is believed to be on a final stage of back-arc spreading. The Shikoku
basin occupying the northern part of the Philippine Sea plate, is an inactive back-arc basin. The major
evidences are magnetic lineations identified by, for example, Watts and Weissel (1975), Kobayashi and
Nakada (1978), and Shih (1980).
In the west of the Izu-Ogasawara arc, back-arc rift system occurs in the northern part of the arc from
Hachijojima Is. to Nishinoshima Is. just behind the volcanic front and in the Mariana trough to the south
of 25°N. The Mariana trough is unfortunately outside the maps of Figs. 1 and 5. In the Mariana trough,
magnetic lineations parallel to the arc were identified (Yamazaki et al., 1991 a). The lineations have been
probably generated by the volcanic activity of back-arc rift system.
In the Japan Sea, there are a number of northeast-southwest trending lineations. The lineations are
inferred to be caused by the products at the center of the opening of the Japan Sea (Isezaki, 1986).
Therefore, the lineations parallel the line of spreading center.
In the Shikoku basin, there are many NWN-SES trending lineations which were initially identified
by Watts and Weissel (1975). These are evidently caused by the products of back-arc spreading of oceanic
plate.
Lineations perpendicular to the arc trend in the back-arc basin
Many lineations are drawn along the Izu-Ogasawara arc from 25°N to 35°N (Fig. 5). Major trend is
west to east, that is, not parallel to the arc trend. Several lineations cross the volcanic front and reaches
the east fore-arc side. Some lineations are evidently related to the Izu-Ogasawara arc Quaternary volcanic
activity because the lineations cross the Quaternary volcanoes.
A similar feature occurs in the Tohoku, where there are several magnetic lineations across the Tohoku
arc.
418 Y. OKUBO et al.
In the Okinawa trough, there are lineations that run through the volcanic arc. They are undoubtedly
caused by igneous rocks of the Ryukyu arc volcanic activity. Besides them, there are a few magnetic
lineations that lie across the Ryukyu arc.
4. Discussions
The main feature in the subduction zone is the deformed magnetic lineations. The depths of
subducting oceanic plate have been estimated from the seismicity (e.g. Kinugasa et al., 1992). The Pacific
plate reaches a depth of 100 km B.S.L. at about 300 km away from the Japan trench. The plate also has
been subducting beneath the Izu-Ogasawara trench. The plate reaches a depth of 100 km B.S.L. at about
150 km away from the Izu-Ogasawara trench. The Philippine Sea plate reaches 20 km B.S.L. at about 100
km away from the Nankai trough and at about 150 km away from the Ryukyu trench. Magnetic layers that
produce magnetic lineations are assumed here to have been subducting with the oceanic plates. The
assumption generates the magnetic layer model of subducting plate expressed in Fig. 6. The Curie point
depth is defined to be a depth where rocks loose their intense magnetization. The existence of magnetic
layers that produce the lineations indicates that ambient temperature ofthe source is below the Curie point.
Therefore, the greatest depth of the source is the minimum depth of the Curie point. The Curie point depths
shown in Fig. 6 are the minimum depths estimated by the greatest depths of the magnetic layers.
Two features in magnetic lineations of the back-arc rifting and spreading system are (1) lineations
parallel to the arc trend, and (2) lineations perpendicular to the arc trend. Origin of the former lineations
is, as many authors suggest, the volcanic activity along the length of the arc as demonstrated in Fig. 7(a).
However, it is not easy to answer a question derived from the second feature of lineations: Why magnetic
lineations occur across the arc trend?
Since several lineations cross the Quaternary volcanoes, the sources at cross parts are at least igneous
rocks. General petrological study from hand samples suggest that magnetic anomalies are mostly caused
by igneous rocks. Therefore, the sources of the entire lineations can be igneous rocks. From the assumption
that the source are igneous rocks, such a model as igneous rocks lie across the arc trend (Fig. 7(b)) can
(a) (b)
450 km 400 300 200 100 0 250 km 200 100 0
Volcanic Front Japan Trench Volcanic Front Izu-Ogasawara Trench
0 0
Tohoku Pacific Plate Izu-Ogasawara Arc Pacific Plate
Overriding Plate urie Isotherm Overriding Plate - Magnetic Layer
100 (Minimum Curie Magnetic Layer 100 Curie Isotherm
Point Depth) (Minimum Curie
Point Depth)
200 200
km km
(C) (d)
350 km 300 200 100 0 200 km 100 0 1 1 1 1 1 1 , 1
Volcanic Front Nankai Trough Volcanic Front Ryukyu Trench
2p Honshu Shikoku 200 Ryukyu Arc Phili Sea 40 Philippine Sea Curie Isotherm pp
km Overriding Plate Curie Isotherm Overriding (Minimum Curie Plate Plate (Minimum Curie Magnetic Layer Plate Point Depth) 100 Point Depth) Magnetic Layer
km
Fig. 6. Subduction model of magnetic layer. (a) Tohoku arc. (b) Izu-Ogasawara arc. (c) Southwest Honshu arc. (d) Ryukyu arc.
The horizontal extension of magnetic layer corresponds to the landward limit of magnetic lineations. The source of the
magnetic layer is assumed to lie in the subducting plate. The Curie isotherm lies beneath the source that produces magnetic
anomalies. Subduction angles of oceanic plates are taken from Kinugasa et al. (1992).
Magnetic Anomalies of Japan and Adjoining Areas 419
(a)
Negaitive Positive Anomalies
Anomalies
~Ob
Q~w
overriding Plate ~,
U (b)
Magnetic
Anomalies Cooled Magma Q~`tr
(Magnetic Source) , plate c overnd~n
Fig. 7. Two magnetic models of back-arc rift system. (a) Magnetic sources parallel to the arc trend. (b) Magnetic sources
perpendicular to the arc trend. The models are based on the assumption that the magnetic lineations are caused by igneous rocks.
be made.
Evidences that the back-arc rift just behind the volcanic front in the northern Izu-Ogasawara arc is
young, have been recently found by Taylor et al. (1990), Urabe and Kusakabe (1990), and Yamazaki et
al. (199lb). In the Okinawa trough, the starting age of crustal separation is estimated to be about 1.9 Ma
by Kimura (1985). The Japan Sea and the Shikoku basin are thought to be on a final stage of back-arc
spreading and now inactive. Therefore, the two types of lineations found in the back-arc spreading regions
can be described as, (1) lineations running parallel with the spreading center of the basins on a rifting or
spreading stage, and (2) lineations across the arc trend in the young back-arc rift region.
As suggested by Tamaki (1985), the two stages are also in different stress field. Elongate magnetic
anomalies related to active faults are found in surrounding regions of the San Andreas fault (Brabb and
Hanna, 1981), the Anatolian fault (Isikara et al., 1985), the Tanna fault (Okubo et al., 1991 b), the Beppu
Bay (Honkura et al., 1994), the Aso volcano (Okubo and Shibuya, 1993). The fault is an exposure of
tectonic movement caused by regional stress field in the crust. Ifthe stress field causes faults perpendicular
to the arc trend, the magnetic anomalies associated with the faults would be produced.
Magnetic sources in the model of Fig. 7 are assumed to be igneous rocks. If compressional or
extensional stress causes a crustal weakness and a magma body penetrates the crustal weakness, difference
in stress field would be a major cause of the two types of volcanic activity.
We wish to express our appreciation to H. Kinoshita and H. Honkura who reviewed our paper and gave helpful
comments.
420 Y. OKUBO et al.
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|
Okubo et al. 1994 Magnetic anomalies of Japan and Adjoining Areas.txt
|
Geochemical Journal, Vol. 25, pp. 335 to 355, 1991
Attainment of solution and gas equilibrium in
geothermal systemsJapanese
HITOSHI CHIBA
Institute for Study of the Earth's Interior, Okayama Univertity,
Misasa, Tottori 682-01, Japan
(Received August 27, 1990; Accepted April 4, 1991)
The geothermal fluids in seven Japanese geothermal systems are tested for attainment of aqueous
and gaseous equilibrium. The pH of fluids in the geothermal reservoir is approximately buffered by the
assemblage K-feldspar-K-mica-quartz. (Na+)/(K+) and (Na+)/ ( activity ratios are ther
modynamically approximated by reactions between albite and K-feldspar, and between albite and anor
thite (or Ca-zeolites), respectively. The (Mg2+)/(K+)2 activity ratio of high temperature geothermal
fluids of Japan can be , represented by the reaction involving Mg-chlorite and K-bearing silicate minerals,
though at lower temperatures other reactions may be responsible. The geothermal fluids are also com
monly saturated with respect to anhydrite and calcite. A small amount of steam loss in the reservoir does
not significantly affect the aqueous composition of the fluids. The partial pressure of CO2 is controlled
by the reaction involving calcite, K-bearing silicate minerals, and albite or Ca-zeolite in geothermal
systems which are not affected by steam loss and dilution. Equilibrium between CH4, C02 and H2 is at
tained at high temperatures but not maintained to lower temperatures in most Japanese geothermal
systems. The H2/H2S ratio is probably equilibrated with Fe-bearing minerals. Gaseous compositions are
very good indicators to identify processes in the geothermal reservoir, such as boiling and dilution. Last
ly, the major aqueous composition and pH of Japanese neutral Na-Cl type geothermal fluid are predic
table if two variables (e.g., temperature and one of the cation activities) are provided.
INTRODUCTION
In the last decade, many geothermal systems
have been assessed and exploited in Japan. A
large number of deep wells have been drilled into
the geothermal reservoirs, and fluids discharged
from them have been analyzed for their
chemical, gaseous and isotopic compositions.
Some of these compositions are now available in
the literature (e.g. Kirishima: Kodama and Naka
jima, 1988; Okuaizu: Nitta et al., 1987). Most
reservoir fluid compositions are Na-Cl dominant
and represent the composition of fluid which in
teracts with reservoir rocks and forms alteration
minerals. Their compositions are likely to be con
trolled by minerals of the rock matrix depending
on the degree of water-rock interaction and also
by the volatiles added to the system from the
magmatic heat source (e.g., Giggenbach, 1984).Studies of reactions between geothermal fluids
and the reservoir mineral assemblage provides
basic information about processes governing
hydrothermal mass transfer in the shallow part
of the earth's crust.
Arnorsson et al. (1983a) examined composi
tions of many Icelandic geothermal waters and
showed that the major element composition is
predictable to as low as 50°C, provided two
parameters, e.g. temperature and chloride con
centration, are given. They concluded that this is
possible due to the attainment of, or close ap
proach to, an overall chemical equilibrium in
geothermal systems. Giggenbach (1980) dis
cussed reactions involving gases in New Zealand
geothermal systems, and concluded that the com
position of fluids reflects close to complete
equilibrium within the system H20, C02, H2S,
NH3, H2, N2 and CH4. The C02 content of the
335
336 H. Chiba
fluid in equilibrium with alteration minerals in
volving calcite and chalcedony is also predicted
by Giggenbach (1981, 1984). In Japanese geother
mal systems, however, the details of reactions
and chemical equilibrium among geothermal
fluids, gases, alteration minerals and reservoir
rocks have not been discussed.
In this study the aqueous speciation and gas
equilibria of neutral Na-Cl type geothermal
fluids from seven Japanese geothermal systems
and some hot spring waters are used to discuss
the controls on the chemical compositions of
geothermal fluid by the reservoir rocks.
DATA SOURCES AND CALCULATIONS
Geothermal well and hot spring data
The locations of geothermal areas and hot
springs examined in this study are shown in Fig.
1. The geothermal wells and hot springs used in
this study are listed in Table 1. The analytical
data and sampling conditions of Japanese
geothermal wells used in this study are given in
the Appendix Table. Mineral assemblages ob
served in drill cores and cuttings have not been
reported in detail for most geothermal systems.
Reported mineral assemblages are briefly sum
marized in Table 2 for some geothermal systems.
The fluid discharged from geothermal wells
are less affected than hot spring waters by boil
ing, mixing with local meteoric water, precipita
tion of minerals or leaching from rock. Thus,
they are better representatives of the fluid
undergoing water-rock interaction at depth.
Complete sets of chemical and physical data, i.e.
chemical compositions of the liquid and steam
phase, discharge enthalpy and sampling condi
tion, are required for the purpose of this study.
Unfortunately, the number of such complete sets
of published data are small, and have been
chosen carefully from the literature. Except for
well It of Okuaizu, wells with high excess en
thalpy discharges are omitted from Table 1 and
subsequent calculations. If the discharge en
thalpy is cited in the literature, it is easy to ex
clude excess enthalpy wells. Even if the discharge
enthalpy is not cited in the literature, a geother140°
130°145'
Nigorikawa
135°
® Hatchobaru0 okuaizu
SNCSYC
Sic• TJC45'
Takigami O0 SC
INC,ISC
STC
irishima 0
30'40°
Kakkonda A
Sumikawa
35°
Fig. 1. Location of geothermal systems considered
in this study. Large circles indicate the location of
geothermal systems, and small dots indicate hot
springs. The symbols used in the following figures are
indicated next to the names of the geothermal
systems.
mal well whose Na/K geothermometer tempera
ture (Fournier, 1979) is far from its silica satura
tion temperature (Arnorsson et al., 1983b) is
excluded because this may indicate the addition
of steam to the discharge within the reservoir.
Geothermal wells in Table 1, except for It of
Okuaizu, are believed to have discharge en
thalpies which agree with the estimated reservoir
temperatures. These fluids can safely be assumed
to exist as a single liquid phase in their reser
voirs. Thus, total discharge compositions (deter
mined from steam fractions) reflect the reservoir
liquid compositions. Okuaizu well It is an exam
ple of an excess enthalpy well and is used only in
gas calculations. The reservoir temperature used
for speciation and gas calculation is estimated by
the chalcedony (< 180°C) (Arnorsson et al.,
1983b), quartz (> 180°C) (Arnorsson et al.,
1983b) or Na/K geothermometer (Fournier,
1979).
Aquifer rock types of geothermal systems in
Table 1 can be roughly grouped into two
categories, volcanic and marine sedimentary
rocks. The host rock can be characterized from
the B/Cl concentration ratio of associated
Solution and gas equilibrium
O V
cu
b0
`nv
0
0 O
0
V
0 CIO
tiG 0
0
U
F.
O
w
G44.1
N
N 3*C4
0 Cd
0
*
x
z*
Uca
U
0
d N
HON
00 00
Cd
0 >4 rA
aN
M
N N
2 2 N_ N N
N N
2 2
00 00
gC." 00 00
vcVi N N
2 2
N Na,, 00w
Q6) Cd
zF.
N
2
00
NM
~I1 N
N
00O~
00
2
2 ~ 2
vi p,
2 '^
kn tn2
N
00
W)
N O N
2 N 2
U x
e.i
z~
z
CO
o
O
zz0
tiZ
CO
o w cnx x
M
2
N00
00
CO
N
O
N Wn
N N " N
2 2
O v~
N
O N
h N
2 2
N N v;
N N
2 2
N M N NO
2
O
N
N 2
00
C' 0\ N
2
00
M Ii
00
2
* *
N
2
00 kn
~ZHx~xHcnv)x~zZrn
4 0.4,
N W N M
E N N N 01 0 00
z W v, x
[t (~ ~, N O QU ~j H E., c0o~MU°M
MocO,=.W UU~~UUU~
t~ ~ N O M 00 to en 00 yj N N yj , , .-. . -~ ..~ -+"U000U,.., UU wh~wwti~x~Z U
Ca*
0
0 x
C30
U y
xxa 0O
m
w
0
v ~
. O p
y 00 1-4
94 V
w V
C ti
atop q~
* * * **
*337
338 H. Chiba
Table 2. Summary of alteration minerals"
Geothermal
SystemQtz Chi Mica Lm/Wa Anh K-feld Cc Py others
Nigorikawa2)
Sumikawa3)
Okuaizu4)
Hatchobaru5)
Takigami6)
Kirishima')+
+
++
+
+
+
++
++
+
+
++
+
+
+
+
++
+
+
++
+
+
+
++
+
+
+sericite
sericite
prehnite
epidote
sericite
prehnite
epidote
Qtz: quartz, Chi: chlorite, Lm/ Wa: laumontitel wairakite, Anh: anhydrite, Cc: calcite, Py: pyrite.
1) The mineral which is explicitly reported to exist at the production level is noted with a "+ " sign.
2) Yoshida (1991)
3) Sakai et al. (1986); Mitsubishi Material Co. (private communication).
4) Nitta et al. (1987).
5) Kyushu Electric Co. (private communication).
6) Hayashi et al. (1988).
7) Kodama and Nakajima (1988).
1000
m100
10
14
0
water0
10 100 1000 ;10000
CI, mg/I
Fig. 2. Relationship between boron and chloride
concentrations of fluids from geothermal wells. Con
centrations are in liquid phase after steam separation.
The lines in the figure stand for molar ratio of B/ Cl.
Data for Nigorikawa and Kakkonda are from Shigeno
and Abe (1987) and Y. Yoshida (private communica
tion), respectively. Symbols are as in Fig. 1.crustal
10He`%-N2magmatic 0 .01 N2
gas (Usu) X
40
+80
+
+,0
+#
+4air
air saturated
.w groundwater
geothermal fluids (Shigeno and Abe, 1983).
Figure 2 is a plot of B against Cl concentrations
in the liquid phase of geothermal well
discharges. Fluids with B / Cl molar ratios be
tween 0.02 and 0.07 (Nigorikawa, Takigami,
Hatchobaru and Okuaizu) discharge from
volcanic reservoir rocks. In Takigami and Hat
chobaru, the geothermal fluids are considered to
be stored in the volcanic rocks (Hayashi et al.,
1988; Manabe and Ejima, 1984). On the other0 20
%-Ar60 80 I.".Ar
Fig. 3. Relative composition of Ar, He and N2 in
geothermal fluids, hot spring waters, mineral spring
waters and volcanic gases. The composition of
magmatic gas is defined by volcanic gases of Mt. Usu,
Hokkaido, Japan (Matsuo et al., 1982). Large plus
signs: volcanic gases of Mt. Usu, small plus signs:
volcanic gases of other volcanoes (Kiyosu, 1985;
Kiyosu and Yoshida, 1988), small dots: gases in
Japanese mineral and hot spring waters (Urabe et al.,
1985), solid squares: fluids in Nigorikawa (Yoshida,
1991), open squares: Okuaizu (Nitta et al., 1987),
solid triangles: Sumikawa (Ueda et al., 1991), open
triangles: Kakkonda geothermal system (Kiyosu and
Yoshida, 1988).
hand, wells in Kirishima, Kakkonda
Sumikawa have B / Cl ratios higher than and
0.07,
Solution and gas equilibrium 339
and are considered to discharge fluids having in
teracted with marine sedimentary rocks. In
Kirishima, the basement rock, Shimanto
Supergroup, is rich in fractures allowing fluid
rock interaction during ascent, though the main
geothermal reservoir is andesitic rocks at
shallower depth (Kodama and Nakajima, 1988).
The aquifer rock types will be discussed later in
relation to the fluid compositions.
The contribution of magmatic gas to the
geothermal systems can be assessed on an Ar
He-N2 diagram (Giggenbach, 1986). Ar and He
are not reactive at hydrothermal conditions. N2
is also not reactive in conditions of Japanese
geothermal systems examined here, since no ap
preciable NH3 is reported. Therefore, these gases
are able to preserve their source signatures after
interaction with reservoir rocks. The gaseous
compositions of Nigorikawa (Yoshida, 1991),
Kakkonda (Kiyosu and Yoshida, 1988),
Sumikawa (Mitsubishi Material Co., private
communication) and Okuaizu (Nitta et al., 1987)
are plotted in Fig. 3 together with those of
Japanese volcanic gases (Kiyosu, 1985; Kiyosu
and Yoshida, 1988) and mineral and hot springs
(Urabe et al., 1985). The magmatic component is
defined in Fig. 3 by the volcanic gas of Mt. Usu
(Matsuo et al., 1982). Figure 3 indicates that the
geothermal gases of Nigorikawa and Okuaizu
have a relatively large contribution of magmatic
gases. In Kakkonda and Sumikawa, geothermal
gases are negligibly affected by a magmatic com
ponent. Some wells in Sumikawa are highly
affected by magmatic component (Ueda et al.,
1991), but they are not used in this study because
of their large excess discharge enthalpy.
Most hot spring waters are affected by boil
ing, mixing of deep fluids with waters of
different origin (e.g. groundwater and steam
heated water), precipitation of minerals from the
water during upflow and/or interaction with the
host rock (e.g., Giggenbach, 1988). Their com
positions do not usually represent the composi
tion of fluid interacting with the rocks in the
reservoir where geothermal fluids are stored. The
chemical composition of over two thousand hot
springs in various Japanese geothermal systemswas compiled by Hirukawa et al. (1977). The hot
springs least affected by shallow processes have
been selected on the basis of the following
criteria: (1) Na/K thermometer temperature
(Fournier, 1979) agrees with chalcedony
(< 180°C) or quartz (> 180°C) saturation tem
perature (Arnorsson et al., 1983b) within 10°C
and (2) the charge balance after aqueous specia
tion calculation is less than 1%. Twenty four of
the over two thousand hot springs cleared the
two criteria and are used in this study. The
speciation calculations were carried out at the
temperatures estimated by the applicable Si02
geothermometer.
For the purpose of comparison, data of
Icelandic and New Zealand geothermal waters
are taken from Arnorsson et al. (1983a) and
Hedenquist (1990), respectively. Hot spring data
in Iceland were selected on the basis of the same
criteria as for Japan. Aqueous speciations were
re-calculated by the code mentioned in the next
section.
Aqueous speciation and gas calculation
The aqueous speciations are calculated using
the code described by Chiba (1990). Forty five
aqueous species are included in the calculation.
It treats only a single liquid phase at tempera
tures from 25 to 300°C. For calculation of a
geothermal well discharge, the steam phase, in
cluding gases, is condensed back to the liquid
phase in proportion to the steam fraction at the
time of sampling. Aqueous speciation of fluid
with excess discharge enthalpy cannot be
calculated by the code used in this study. The
dissociation constants of aqueous species are
adopted from the thermodynamic data base of
SOLVEQ (Reed, 1982). The thermodynamic
data of minerals are from Helgeson et al. (1978).
The calculation of gaseous species follows
the method of Giggenbach (1980). Fugacity
coefficients of gases are assumed to be unity in
the same manner as Giggenbach (1980). C02
mineral equilibrium constants were generated by
SUPCRT (Helgeson et al., 1978) using the 1981
data base. This results in stability relationships
of Ca-Al-bearing minerals that are slightly
340 H. Chiba
different from the original figures in Giggenbach
(1984).
RESULTS OF AQUEOUS SPECIATION AND
GAS CALCULATIONS
Aqueous species
In Figs. 4 and 5, cation/proton activity
ratios are plotted against reservoir temperatures.
Except for the (Mg")/(H')' ratio, other ratios
show simple patterns against reservoir tempera
ture. This suggests that these ratios are controll
x
z
rn
07
6
5
000%
0 • •0
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_ K-feld~rF
(=rnlq"
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9
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least squares
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8
6
Fig. 4. The
tion/proton
discharges
recalculate
tion /proton
The dashed
Symbols are as in Fig. 1.4
150 200 250 300
Temperature, °C
temperature dependence of ca
activity ratios of geothermal well
in Japan. The lines in the figures are
d temperature dependences of ca
ratios in Icelandic geothermal waters.
curve in Fig. 4B represents reaction (3).150 200 250 300
Temperature, °C
Fig. 5. The temperature dependence of ca
tion/proton activity ratios of geothermal well
discharges in Japan. The solid curves in the figures are
recalculated temperature dependences of ca
tion/proton ratios in Icelandic geothermal waters.
The lower curve in Fig. SA is a least squares fit of
some of the plotted data (see text). Symbols are as in
Fig. 1.
ed by mineral buffer systems in the reservoir.
As inferred from Fig. 2, the aquifer rock
types of Nigorikawa, Okuaizu, Hatchobaru and
Takigami geothermal systems are volcanic rocks.
Marine sedimentary (or metasedimentary) rocks
are judged to be the aquifer rocks of Kakkonda,
Sumikawa and Kirishima geothermal systems.
Figs. 4 and 5 indicate that the aquifer rock type
does not systematically affect the major element
fluid compositions. The same phenomena were
observed in Icelandic thermal waters (Arnorsson
et al., 1983a) and in water/rock experiments
(Kacandes and Grandstaff, 1989).
Arnorsson et al. (1983a) showed that ca
tion/proton and cation/cation activity ratios of
Icelandic thermal waters follow simple patterns
against reservoir temperatures as low as 50°C.
The patterns are considered to be the results of
silicate mineral buffer systems, which may vary
with temperature but have smooth transitions
Solution and gas equilibrium 341
because of small differences in the A G of
minerals. However, the patterns given by them
cannot be directly compared with the present
results, because some of their thermodynamic
data used for speciation are different from those
used in this study. The curves in Figs. 4 and 5
were obtained by least squares fits of the
.recalculated speciations of Icelandic samples.
Except for (Mg")/(H')', most cation/proton
ratios scatter around the curves obtained from
Icelandic samples. This suggests that the ca
tion/proton ratios in Japanese geothermal
systems may be controlled by silicate mineral
buffer systems similar to Icelandic geothermal
systems. Detailed discussions about silicate
mineral buffer systems follow examination of the
individual geothermal systems, particularly on
effects disturbing the aqueous and gaseous com
position, such as boiling in the reservoir and mix
ing with waters of different origin. The hot
spring waters are not plotted in Figs. 4 and 5 as
they have a large scatter. Though they were very
carefully selected from the literature, their
erratic results indicate that the waters discharged
from Japanese hot springs have been affected by
processes such as boiling, mixing with shallow
water and/or reaction at low temperature with
minerals not accounted for by the curves in Figs.
4 and 5 before they reach the surface. Therefore,
they are not good representatives of fluid interac
ting with rock at reservoir depths.-10
-15
~, -20 Y
° -25
-30
-35 jr lib•v
o.~
000 0
of0
0
P s3 100"P
a
Gaseous species
Gaseous species are treated in the manner de
scribed by Giggenbach (1980). He also con
sidered the effect of steam loss and gain, and the
attainment of equilibrium in the carbon and
sulfur gas systems is discussed following his inter
pretive framework.
Carbon dioxide and methane are carbon
bearing gases whose concentrations are often
analyzed in geothermal studies. Between these
two gases, the following equilibrium often ex
ists:
CO2+4H2=2H2O+CH4. (1)
The analytical equilibrium constants for this100 150 200 250 300
Temperature, °C
Fig. 6. Plot of analytical log K for reaction (1) ver
sus reservoir temperature, illustrating the effect of
steam loss and gain on C02-H2-CH4 gas equilibrium
(Giggenbach, 1980). The arrow shows an example of a
boiling path for a fluid whose gas composition is in
equilibrium for reaction (1) at 245°C. Symbols are as
in Fig. 1.
N
x
01
0
-1
-2
-3A
lot
'Mite
P;-We--O
Am
e-
0 !
~b oo~
100 200 300
Temperature, °C
Fig. 7. Log H2/H2S of total discharge versus reser
voir temperature. The curves are for pyrite-Fe-AI
silicate, pyrite-magnetite and pyrite pyrrhotite buffers
(Giggenbach, 1980). The arrow shows an example of a
boiling path for a fluid whose composition is buffered
by the pyrite-Fe-Al-silicate reaction (2) at 245°C.
Symbols are as in Fig. 1.
reaction (log Kc") of gases from geothermal
wells are plotted in Fig. 6. The curve at the
center corresponds to the equilibrium of reaction
(1) for all species dissolved in a single liquid
phase. The arrow originating from equilibrium
log Kc" at 245°C is an example of a predicted
boiling path, which is calculated assuming
adiabatic continuous vapor loss. Gaseous com
positions of Japanese geothermal systems scatter
widely around the equilibrium curve. Samples
lying in the region between equilibrium and
342 H. Chiba
equilibrium vapor indicate a fluid which gained
excess steam in the reservoir (Giggenbach, 1980).
The fluid of the Okuaizu system has clearly gain
ed excess steam. In contrast, all fluids from the
Kirishima geothermal system plot below the
equilibrium curve, and on a trend indicating
steam loss from a fluid that may have begun boil
ing at about 245'C; the degree of steam lost is in
dicated by the contours of steam fraction (Gig
genbach, 1980). Equilibrium of reaction (1) ap
pears to be attained in the Kakkonda and
Nigorikawa geothermal systems.
Ratios of H2/H2S mole fraction in total
discharge are plotted against reservoir tempera
tures in Fig. 7. Equilibrium H2/H2S ratios for
pyrite-pyrrhotite, pyrite-magnetite and pyrite
Fe-Al-silicate coexistence given by Giggenbach
(1980) are also shown by the curves in Fig. 7.
The arrow originating from the equilibrium
log(H2/H2S) value of pyrite-Fe-Al-silicate at
245'C indicates an example of an adiabatic boil
ing path. Data for Japanese geothermal wells
scatter widely, though some patterns are visible.
The H2/H2S ratios of some geothermal well
discharges must be influenced by steam loss in
the reservoir. The aqueous solubility of H2 is
much lower than that of H2S, because the gas
distribution coefficient of H2 between vapor and
liquid is greater than that of H2S (Giggenbach,
1980). Therefore, steam (and gas) loss from the
reservoir lowers the H2/H2S ratio, and the
downward scatter from the buffer systems shown
in Fig. 7 may be the result of steam loss from the
reservoir fluid. Fluids from the Kirishima
geothermal system again show a simple vapor
loss trend as for the carbon-bearing gases (Fig.
6). Fluid samples which suggest steam loss in
Fig. 6 (higher temperature wells in Takigami and
one well in Sumikawa) also plot below any buffer
ing systems in Fig. 7, supporting steam loss in
these reservoirs. In the Kakkonda system,
H2/H2S ratios of fluids may be controlled by
pyrite-magnetite in some wells and pyrite-Fe-Al
silicate buffers in other wells. The latter buffer
system was empirically deduced by Giggenbach
(1980) according to the following reaction: pyrite + H2+ H20= FeO(silicate) + 2H2S, (2)
where FeO (silicate) means Fe" in Al-bearing
silicate mineral, e.g. chlorite. The H2/H2S ratio
of one well in Sumikawa may also be accounted
for by reaction (2). The pyrite-pyrrhotite buffer
system cannot be totally ruled out for Japanese
geothermal systems since the vapor loss sug
gested for some wells makes a clear determina
tion of the actual buffer system difficult.
CHARACTERISTICS OF INDIVIDUAL
GEOTHERMAL SYSTEMS
Reactions among aqueous species and silicate
minerals appear close to equilibria in Japanese
geothermal systems (Figs. 4 and 5). However,
boiling and steam gain in the geothermal reser
voir affect the gaseous compositions of some
systems and cause the gaseous compositions to
shift from equilibrium states (Figs. 6 and 7).
These results probably reflect the processes occur
ring in the reservoir of individual geothermal
systems. In other words, partial equilibrium is
frequently attained depending on specific condi
tions in each geothermal system. Before discuss
ing the details of reactions controlling the com
positions of geothermal fluids, we must
recognize the processes influencing the fluid com
position of individual geothermal systems. In
this section, based on the degree of partial
equilibrium attained in an individual system, the
characteristics of each geothermal system will be
briefly discussed.
Nigorikawa geothermal system
The C02 content of fluids in Nigorikawa is
much higher than in other systems (Appendix
Table). Limestone is one of the components of
the geothermal reservoir (Sato, 1988). A possible
explanation of the high CO2 flux is attack of
acidic gas of magmatic origin on limestone to
produce CO2. Despite the addition of a large
amount of CO2 gas to the system, the aqueous
species appear to have approached equilibrium
with silicate minerals, except for (Ca2+)/(H+)2
(Figs. 4 and 5). The high flux of CO2 influences
Solution and gas equilibrium 343
the concentration of aqueous Ca species, since a
large amount of CO2 in the liquid phase will
cause calcium in solution to precipitate as
calcite. The deficiency of Ca2+ can be seen in
Fig. 4C. However, activities of other cations and
pH are not significantly affected by the high C02
flux.
The reaction between CH4, C02 and H2 is
very close to equilibrium, indicating that exten
sive boiling in the geothermal reservoir does not
take place. The contribution of magmatic gases
indicated from Fig. 3 in the Nigorikawa system
does not affect the attainment of equilibrium for
reaction (1). Therefore, fluid discharged from
the Nigorikawa system is representative of fluids
stored in the geothermal reservoir, though the
Ca 21 is slightly influenced by the high C02 flux l empirical curves determined by Icelandic ther
mal waters, except for (Mg2+)/(H+)2 of one fluid
sample. The aqueous species are likely controll
ed by silicate mineral assemblages. Gaseous reac
tions are slightly out of equilibrium, indicating a
small steam loss in the fluid from well S-4.
Okuaizu geothermal system
The one sample from Okuaizu is used in this
study as an example of a fluid with an excess
discharge enthalpy. Though the aqueous com
positions are not plotted in Figs. 4 and 5, they
are far from the empirical curves. The aqueous
composition is strongly influenced by the gain of
steam in the geothermal reservoir, as indicated in
Fig. 6. This sample will be omitted from the
discussion of aqueous equilibrium.
Kakkonda geothermal system
The cation/proton ratios of Kakkonda are
very close to the pattern of Icelandic geothermal
waters, except for (Mg2+)/(H+)2 (Figs. 4 and 5).
The (Mg2+)/(H+)2 ratio is controlled by a reac
tion which includes Mg-chlorite, as discussed
later. Thus, the aqueous composition of fluid in
Kakkonda is closely controlled by the silicate
mineral buffers. Most samples of fluid in Kak
konda are in or close to equilibrium with the
gaseous reaction involving CH4, CO2 and H2,
though one well indicates a slight steam gain
(Fig. 6). The H2/H2S ratio of some samples ap
pears to be controlled by the reaction involving
magnetite and pyrite, and others by the reaction
of Fe-Al-silicate and pyrite, as mentioned earlier
(Fig. 7). Unfortunately, these buffer systems can
not be confirmed by field observation since de
tailed study of Fe-bearing minerals has yet to be
reported. Figures 4 to 7 suggest that fluids in the
Kakkonda system are in full equilibrium with the
reservoir rock.
Sumikawa geothermal system
. The Sumikawa geothermal system is located
about 20 km from the Kakkonda geothermal
system. The salinity of fluid (0.006 mole Cl
kg/H20) is the lowest among fluids examined
here. The cation/proton ratios are close to theHatchobaru geothermal system
All aqueous species in the Hatchobaru
system are close to equilibrium. Since the gas
data required for calculating gaseous equilibria
are not available, boiling and/or steam gain are
not assessed in Figs. 6 and 7. However, Pco2 ap
pears to be closely buffered by the mineral
assemblage, as discussed later. This means that
samples used in this study are not likely affected
by boiling or steam gain. The fluid in the Hat
chobaru system, including gases, is in full
equilibrium with the reservoir rocks.
Takigami geothermal system
A plot of Cl concentrations vs. discharge en
thalpies for the Takigami system (Takenaka and
Furuya, 1991) indicates a strong dilution trend
from the parent fluid (TT-14 well) towards
shallow water whose Cl concentration can be
represented by local hot spring waters. Samples
from wells NE-2 and NE-3 (not plotted in
Takenaka and Furuya, 1991) also lie on the same
dilution line originating from the fluid of TT-14;
they are the most diluted well discharges of
Takigami. The aqueous compositions plot close
to the empirical curves with trends crossing the
empirical curves (Figs. 4 and 5). These trends
might be characteristic of fluids diluted by
shallow water (e.g., Hedenquist, 1990), though
344 H. Chiba
the cation/cation activity ratios are close to
equilibria with alteration minerals, as discussed
later. Compositions of carbon-bearing gases
deviate from equilibrium at high temperature
and approach equilibrium at low temperature.
This erratic result may indicate that reaction
among CH4, CO2 and H2 does not take place in
the dilution process at Takigami and that the
composition of carbon-bearing gases in the
parent fluid which had lost steam in the reservoir
is preserved in the diluted fluids. The composi
tion of fluids in the Takigami geothermal system
is affected by progressive dilution across the
system.
Kirishima geothermal system
All cation/proton activity ratios in this
system plot near the empirical curves, but scatter
wider than in other geothermal systems (Figs. 4
and 5). All gas data indicate steam loss in the
reservoir (Figs. 6 and 7), with boiling beginning
at about 245'C. The fluid with the largest steam
loss, as indicated in Fig. 6, shows the largest
deviation from equilibrium conditions in Figs. 4
and 5. This suggests that aqueous compositions
are slightly affected by steam loss in the reser
voir. Among the geothermal systems examined
here, Kirishima appears to be representative of a
system influenced by steam loss.
DISCUSSION
Effects of boiling and dilution on aqueous and
gas compositions
The solute-mineral, gas-gas and gas-mineral
equilibria discussed earlier appear to be most
closely attained in the Kakkonda and Hat
chobaru geothermal systems, whose reservoir
temperatures cover the range of the studied
geothermal systems, except for low temperature
Takigami wells. In Japanese geothermal
systems, gaseous compositions scatter wider
around empirical or theoretical equilibrium
values than aqueous compositions, suggesting
higher sensitivities of gaseous species for boiling
in the reservoir and/or mixing.
Figures 6 and 7 indicate that the fluids in theKirishima geothermal system are influenced by
boiling, as previously mentioned. According to
Fig. 6, the amount of steam lost from the reser
voir is as high as 10% at Kirishima, assuming
that reaction (1) controlled the initial gas com
position. Reed and Spycher (1985) calculated the
boiling effects on a Broadlands-like fluid in a
fully equilibrated system, in which instantaneous
equilibrium is always attained among minerals,
gases and aqueous species. According to them,
ten percent isoenthalpic boiling of an initially
278'C fluid to 243'C causes a pH increase of
about 0.5 unit, but does not significantly change
simple cation activities. Thus, boiling causes the
cation/proton ratios to deviate to higher values
(Figs. 4 and 5). The results in Figs. 4, 5, 6 and 7
support boiling and steam loss in the reservoir,
with the data from Kirishima roughly agreeing
with the temperature-dependent curve if they are
corrected for steam loss and pH increase. Steam
and gas loss clearly affects not only the gaseous
compositions but also aqueous compositions to
a lesser extent.
The effects of dilution by water of shallow
origin can be seen in the Takigami geothermal
system. In considering reactive elements (e.g. Na
and K) versus a mobile element (Cl), fluid com
positions should plot on a mixing line, if there is
no reaction between aqueous species and
minerals after the mixing. However, this is not
the case for the Takigami geothermal system, as
fluid compositions deviate from simple mixing
lines (Takenaka and Furuya, 1991). The devia
tions of fluid compositions from mixing lines sug
gest two possibilities: (i) the mixing is not a sim
ple two endmember mixing or (ii) reactions
between aqueous species and the reservoir rock
take place to some extent after mixing. In the
Takigami geothermal system, reservoir tempera
tures calculated by the Na-K-Ca geother
mometer agree well with those of the Si02
geothermometers (Appendix Table), and Ca"
and SO4 concentrations of fluids are saturated
with respect to anhydrite (Hayashi et al., 1988).
Therefore, the second possibility is favored for
the Takigami geothermal system, though
equilibria between aqueous species and rock are
Solution and gas equilibrium 345
a.
a.9
8
7
6
5
4
9
8
7
6
5JapanA
Q-01 O'o~ 0.01
pO
chalcedonyquartz
lvt~ba ~9--; 000,
Qr ..
Iceland
chalcedonya•
quartzB
0.01
9.1
0 100 200
Temperature, °C
dependence300
Temperature
H20.Fig. 8. A: of pH in
Japanese thermal waters. Lines indicate the tempera
ture dependence of pH when pH is buffered by the K
feldspar-K-mica-quartz (or chalcedony at less than
200°C) assemblage at a Na+K concentration of 0.1
and 0.01 mole/kg Symbols are as in Fig. 1.
B: Temperature dependence of pH of Icelandic ther
mal waters. Large circles indicate well discharges.
Small dots represent hot spring waters.
not complete, as indicated by the scatter of ca
tion/proton activity ratios (Figs. 4 and 5). The
gas composition also appears to be out of
equilibrium (Figs. 6 and 7).
Control of pH of Japanese thermal waters
mole/kg H20
K-feldspar + H + The pH values at reservoir temperatures are
calculated for Japanese thermal waters and are
plotted in Fig. 8A. The pH of Icelandic thermal
waters are also plotted in Fig. 8B for com
parison. The lines in Fig. 8 indicate a tempera
ture dependence of pH, assuming (i) concentra
tion of E (Na + K) indicated in
Fig. 8, (ii) the relation between activity ratio of
Na+ / K+ and temperature applies (Arnorsson et
al., 1983a) and (iii) the following reaction:
1.5
= 0.5 K-mica + 3 quartz(or chalcedony) + K+ .
(3)Chalcedony saturation was assumed at less than
200°C. The assumed Na+/K+ activity ratio is
very close to a fluid which is in equilibrium with
low albite and microcline (Arnorsson et al.,
1983a).
(Na + K)
H20,
mole/ H20
H20
H20, Geothermal wells in Japan have pH values
around the line of concentration of
0.01 mole/kg except for wells in
Nigorikawa. As the E (Na + K) concentrations
of fluids from these wells range from 0.009 to
0.05, the reservoir pH values appear to be con
trolled by reaction (3). The reservoir pH of
Nigorikawa is lowest among all wells in this
study. Wells in Nigorikawa discharge fluid with
the highest concentration of E (Na + K) and E
CO2 in Japan. The E CO2 concentration is as
high as 0.14 kg and E (Na + K) con
centration is 0.21 mole/kg on a total
discharge basis, with the high CO2 probably
related to the limestone-bearing basement rocks.
The reservoir pH of Nigorikawa is close to the
pH predicted by the silicate mineral buffer
system for the measured E (Na + K) concentra
tions of 0.21 mole/kg indicating the high
E CO2 concentration does not affect the reser
voir pH. Therefore, the reservoir pH of fluids
discharged from Japanese geothermal wells are
approximately buffered by a K-bearing silicate
mineral assemblage.
C02 to
(Na + K)
Na + K The pH values of Japanese hot spring waters,
especially those of high temperature, show a
much larger scatter around the pH mineral
buffer than Icelandic thermal waters. This may
result from the loss of a small amount of
the vapor during boiling to the surface. On the
other hand, the pH values of Icelandic hot
spring waters follow the theoretical temperature
concentration. dependence at low E
This is expected because most Icelandic hot
concenspring waters have uniformly low
trations of around 0.001 mole/kg H2O. Also,
Icelandic well discharges have pH values close to
the pH predicted for their E (Na + K) concentra
tion, assuming a K-bearing silicate mineral
buffer.
346 H. Chiba
Control of aqueous species by silicate mineral
buffers
Saturation of silica minerals, such as quartz
or chalcedony, is implicitly assumed in this
study, and is used to estimate the reservoir tem
peratures of geothermal fluids. The relatively
fast reaction rate between silica minerals and
fluid at geothermal temperature is widely ac
cepted (e.g., Ellis and Mahon, 1977). Also, there
is little silica precipitated during the rapid ascent
of fluid from the bottom of the well to the sur
face. Furthermore, the silica temperatures of
most fluids used in this study agree well with
Na/K temperatures (Appendix Table), suppor
ting the assumption of silica saturation at reser
voir temperatures.
The saturation state of individual Al-bearing
silicate minerals is not easily determined from
the component aqueous species because the ac
curacies of Al analyses in geothermal fluids and
the dissociation constants of Al(OH)n3-n>+
species are frequently not adequate for such
evaluation (Chiba, 1990). Some examples of in
dividual mineral saturations are discussed for
geothermal well discharges by Chiba (1990). Fur
ther investigations about the saturation state of
individual Al-bearing silicate minerals are not
considered in this paper, since the quality of Al
analyses for the geothermal fluid samples studied
here is uneven.
However, we can investigate the reactions
controlling the relative abundance of aqueous
species using cation/cation activity ratios. In
Fig. 9, cation/cation activity ratios of Japanese
geothermal fluids are plotted against reservoir
temperatures. Both (Na+) / (K+) and
(Na+) / (Ca2+) show simple patterns against
reservoir temperature and agree with the em
pirical curve obtained from Icelandic thermal
waters.
The dashed curve in Fig. 9A, calculated from
the thermodynamic data of Helgeson et al.
(1978), is the temperature dependence of the
(Na+)/(K+) ratio of the following reaction:
albite + K+ = K-feldspar + Na+. (4)
The close agreement of Japanese geothermal2.0
Y 1.5
z
0 1.0
0.5
c 1
z
0 0
-1
1 N
Y
0 N
0)
~ -1
s
-2
-3+
150A
®r~
rr~lo
I
B
i a.---------- i'
Anorsson et al. (1983)
I
Jb, Ac ~ele ~tC
O
O
9L*: i9i
Aot0
41e
Fig. 9. The temperature f
tion activity ratios
Solid curves marke
recalculated tempe f
tion/proton ratios 'The dashed curve i p
dependence of (Na+)/(K ff y o
( ). )
Dashed curve in Fig. 9B corresponds to combinations
of reactions (5) at <250 and (6) at > 250°C. The other
two curves correspond to reaction (8); the upper at
chalcedony and the lower at quartz saturation. Sym
bols are as in Fig. 1.200 250 300
Temperature, 00
dependence o cation/ca
of Japanese geothermal fluids.
d Arnorsson et al. (1983) are
rature dependences o ca
of Icelandic geothermal waters.
n Fig. 9A indicates tem erature
+ bu ered b reacti n 4
fluids with the dashed curve indicates the
(Na+) / (K+) ratio can be represented by the reac
tion of low albite and K-feldspar, as pointed out
in other geothermal systems (e.g. Arnorsson et
al., 1983a; Giggenbach, 1988).
The (Na+ / (Ca2+) ratios (Fig. 9B) show
strong correlations with temperature, except for
the Nigorikawa geothermal system, and also
Solution and gas equilibrium 347
agree well with the empirical curve obtained by
Icelandic thermal waters. The scatter of
(Na+)/ (Ca2+) ratios around the curve is much
smaller than that for Icelandic thermal waters.
The Nigorikawa (Na+) / (Ca2+) ratios are
distinctly high compared to those of other
geothermal systems. As (Na+)/ (H+) ratios of
Nigorikawa have the same pattern as those of
other geothermal systems, the deviation of
(Na')/1(_Ca') is due to low (Ca")/(H')'. The
low (Ca2+)/(H+)2 is caused by the high C02 flux
in Nigorikawa system, as discussed previously.
The dashed curve in Fig. 9B is a combination of
the following two reactions:
2 albite + Ca 2+ +4H20 =
laumontite + 2Na+ + 2 quartz (<2501Q, (5)
and
2 albite + Ca' + 2H20
wairakite + 2Na+ + 2 quartz (>250'C). (6)
The dotted curve corresponds to the reaction:
2 albite + Ca2+ = anorthite + 2Na + + 4 quartz.
(7)
Above 200°C, the (Na+)/ (Ca2+) ratios of
Japanese geothermal fluids plot close to the dot
ted curve, indicating the (Na+)/ (Ca2+) ratio of
geothermal fluids can be represented by reaction
(7). Reaction (6) also results in similar
(Na+)/ (Ca +) ratios to reaction (7) above
250°C.
Other information about aqueous species
relating to equilibrium with alteration minerals
can be obtained from the Na-K-Mg triangular
diagram (Giggenbach, 1988). The chemical com
positions of geothermal wells are plotted in this
diagram (Fig. 10), except for Kirishima wells,
where Mg concentrations were not reported.
Also plotted as small dots are Icelandic thermal
waters used to define the curves on Figs. 5, 6 and
9. According to Giggenbach (1988), when full
equilibrium is attained between geothermal fluid
and alteration minerals, reaction (4) and the
following reactions control Na, K and Mg con
centrations:K/100%-Na 60
40
rk:Na/1000
full equilibrium
a ° $° • o
Q • °
immature waters
iMg
Fig. 10. p g
trations of Japanese geothermal fluids and Icelandic
thermal waters. Small dots indicate Icelandic thermal
waters. Large symbols are as in Fig. 1. The region of
immature waters corresponds to the composition of
waters related to rock dissolution without any control
by mineral equilibria (Giggenbach, 1988).0 20 40 60 80 100
-Mg
Relationshi between Na, K and M concen
0.8 K-mica + 0.2 Mg-chlorite + 5.4 silica + 2 K+
=2 .8 K-feldspar + 1.6 water + Mg2+ . (8)
As indicated from Fig. 10, geothermal fluids in
five Japanese geothermal systems (Nigorikawa,
Kakkonda, Sumikawa, Okuaizu and Hat
chobaru) are at or very close to full equilibrium.
However, geothermal fluids of Takigami and
one sample from Sumikawa are in partial
equilibrium in terms of combination of reactions
(4) and (8). Also, about half of Icelandic thermal
waters are far from full equilibrium. Figs. 5A,
9C and 10 suggest that reaction (8) i-s not attain
ed in fluids of Takigami, Sumikawa and some
Icelandic geothermal systems.
In Fig. 9C, two curves are shown in addition
to the empirical curve obtained by. Icelandic ther
mal waters. These curves correspond to reaction
(8); the upper curve uses chalcedony (Giggen
bach, 1988) and the lower one uses quartz as the
silica phase. The fluids, which are indicated to be
in, or close to, full equilibrium (Fig. 10), plot
close to the lower curve, suggesting that their
(Mg2+)/(K+)2 ratios are controlled by reaction
(8) in the presence of quartz. In Fig. 5A an alter
native curve can be drawn by a least squares fit,
connecting fluids in full equilibrium (Fig. 10)
(i.e., all wells of Hatchobaru, Nigorikawa and
348 H. Chiba
-4
-6
-8
a
o -10
-12
-14
-16
-6
-8
a
rn 10 0
-12
-14
-16
-6
-8
a
rn 10 0
-12
-14
-16
0 100 200 300
Temperature, °C
Fig. 11. Plot of logarithm of activity product (log
Q) versus reservoir temperature. Solid circles stand
for activity products of Ca" and SO,', in geothermal
fluids, and open circles for Ca" and COj-. Small and
large circles indicate hot spring and geothermal well
samples, respectively. The upper line is the solubility
product of anhydrite and the lower is calcite. From
top to the bottom, figures are for Japanese, Icelandic
and Broadlands, N. Z., thermal waters, respectively.A supersaturated
r
undersaturated 0
Japan
II I
Bsupersaturated
hte
o 0
undersaturated
Iceland
C supersaturated
N,
0 e0 0 ,a
0
undersaturated
Broadlands, N.Z0C
Kakkonda, and one in Sumikawa). The new
curve (Fig. 5A) may represent the (Mg2+)/(H+)2
ratio when the fluid is buffered by reaction (8) .
The Icelandic empirical curve may correspond to
another reaction applicable at lower tempera
tures, since many Icelandic thermal waters at
temperatures lower than 200°C plot very close to
this curve.
In Fig. 4B the dashed line indicates the tem
perature dependence of the (K+)/ (H+) ratio of
reaction (3), which was calculated from thermodynamic data (Helgeson et al., 1978), assum
ing that activities of all solid phases are unity.
The close agreement of the calculated (K+)/(H+)
ratios of geothermal fluids to the dashed line at
high temperature suggests that the relative abun
dances of K+ and H+ ions are controlled by reac
tion (3), as expected from the role it plays as a
pH buffer of geothermal fluids.
Saturations with respect to anhydrite and calcite
Activity products, (Ca2+)(SO4-) and (Ca2+)
(C02 3 ), are plotted (as log Q) against reservoir
temperature in Fig. 11. The upper line of each
figure indicates the solubility product of
anhydrite and the lower is calcite. All Japanese
geothermal wells are approximately saturated
with respect to anhydrite (Fig. 1 1A). Hot spring
waters are slightly undersaturated, but are much
closer to saturation than Icelandic hot springs
and low temperature geothermal fluids. Geother
mal fluids in Broadlands, New Zealand, are ob
viously undersaturated with respect to
anhydrite, even at high temperature.
Sakai and Matsubaya (1974) categorized
Japanese thermal waters into four types based
on their stable isotopic ratios. One is Green-tuff
type thermal water, which leaches fossil Miocene
marine sulfates that precipitated contem
poraneously with the Green tuff formation.
Most of the Japanese hot springs studied here
are located near the Green tuff formation. In
these cases, approach to anhydrite saturation of
Japanese hot spring waters may be due to the
leaching of fossil marine anhydrite. In some
geothermal systems, however, there is no indica
tion of the influence of fossil seawater on the
composition of geothermal fluid. For example,
in Takigami and Hatchobaru, the B / Cl ratios
(Fig. 2) suggest that the geothermal reservoirs
consist of volcanic rocks. Anhydrite is abundant
in veins of altered rocks at Takigami, and the
aqueous sulfate concentration is controlled by
solubility of anhydrite (Hayashi et al., 1988).
Anhydrite is also found in veins in all of the
other geothermal systems studied here (Manabe
and Ejima, 1984; Yoshida, 1990; Mitsubishi
Material Co., private communication; Kodama
Solution and gas equilibrium 349
and Nakajima, 1988). Therefore, the sulfate con
centrations of geothermal fluids must be controll
ed by anhydrite in the reservoir, similar to that in
Takigami. The common attainment of anhydrite
saturation is a significant characteristic of
Japanese geothermal systems when compared to
systems in Iceland and Broadlands, N.Z. The re
maining question is why anhydrite commonly
forms in Japanese geothermal systems. A stable
isotopic study of anhydrite and dissolved sulfate
may provide the clues to understanding the
widespread presence of anhydrite in Japanese
geothermal systems.
All the fluids discharged from Japanese
geothermal wells in this study are approximately 'saturated with respect to calcite, as expected
from its common presence in drilling cores and
cuttings of geothermal wells. About half the
Japanese hot spring waters are super-saturated
with respect to calcite, whereas Icelandic hot spr
ing waters are just saturated. The surface temper
atures of hot springs whose waters are super
saturated with respect to calcite are close to boil
ing. If a small amount of vapor is lost by surface
or subsurface boiling, much C02 escapes from
the liquid to steam phase because of the large gas
distribution coefficient (Giggenbach, 1980). This
removal of CO2 from the hot spring before it is
sampled causes an increase in pH and C03-,
which results in an increase in the calculated
(Ca2+)(CO3-) activity product. The calculation in
dicates qualitatively that the observed super
saturation of hot spring waters with respect to
calcite results from C02 loss, probably due to
boiling.
Anhydrite and calcite might control gas
equilibrium through the following reaction:
calcite + H2S + H20(1) = anhydrite + CH4. (9)
The logarithm of equilibrium constants, i.e. log
PCH4/PH2s are almost temperature independent at
around -4.9 (Giggenbach, 1980). Log
PCH4/Px2s values of Japanese geothermal well
discharges range from -0.9 to 1.9, indicating
that equilibria are attained among Ca2+, C03
and SO4 ions only in the aqueous phase, and
that gases are far from equilibrium involvinga9
CD 0-2
-1
0
1
2
3 prehnite
i
0100 grossuiar
v
Cap
pyrophyllite
Fig. 12. c
voir temp
minerals f gg
The solid curve marked (C) indicates the temperature
dependence of Pco1 buffered by reactions (9) and (10).
The curve marked (Q) is for reactions (10) and (11) or
(12). Symbols are as in Fig. 1.100 150 200 250 300
Temperature, °C
o2 (bars) and reserRelationship between P
erature. Stability fields of Ca-bearing silicate
recalculated a ter Gi enbach (1980). are
anhydrite and calcite. The equilibrium of reac
tion (9) is also not attained in the geothermal
systems studied by Giggenbach (1980).
Therefore, the reaction rate among anhydrite,
calcite and carbon and sulfur-bearing gases
must be very sluggish.
Partial pressure of CO2
Partial pressures of CO2 are plotted against
reservoir temperatures in Fig. 12. The stability
relations among Ca-bearing minerals are from
Giggenbach (1984), and are drawn assuming the
presence of calcite and chalcedony (i.e. about
10% super-saturated with respect to quartz).
Two solid curves calculated from ther
modynamic data (Helgeson et al., 1978) repre
sent the potential C02 buffer systems involving
K-feldspar and K-layer silicate in a fully
equilibrated system. The lower solid curve mark
ed (C) is a combination of the following two reac
tions:
laumontite + K-feldspar +C02=
calcite + K-mica + 4 chalcedony
+ 3H20 (<250°C), and (10)
wairakite + K-feldspar + CO2 =
calcite + K-mica + 4 chalcedony
+H2O (>250°C). (11)
350 H. Chiba
The upper curve (Q) assumes the presence of
quartz instead of chalcedony. The corre
sponding reactions are as follows:
laumontite + K-feldspar +C02=
calcite + K-mica + 4 quartz
+3H20(< 250'C), and (12)
wairakite + K-feldspar + CO2 =
calcite + K-mica + 4 quartz
+H2O (>250°C). (13)
The following reaction also gives similar Pco2
values to reaction (13) between 250 and 300°C:
anorthite + K-feldspar + C02+ H20=
calcite + K-mica + 2 quartz. (14)
The geothermal fluid of Okuaizu, which has ex
cess enthalpy discharges, has a very high Pco2, as expected from the gain of excess steam. Wells in
Nigorikawa discharge fluids of high C02 con
tent. Fluids from Kirishima, which show a
strong boiling trend in Figs. 6 and 7, are also
best explained by C02 loss accompanying steam
loss. The near horizontal distribution of Pco2 in the Takigami system (Fig. 12) can be accounted
for by dilution of high temperature fluid by a
cooler water, with boiling and gas loss thus being
quenched.
Only the fluids from Hatchobaru plot on the
solid curve (Q). Since temperatures calculated by
the quartz geothermometer for fluids from the
Hatchobaru system agree with Na/K tempera
tures, the prevailing silica phase is expected to be
quartz. Thus, the Pco2 of Hatchobaru wells is
probably controlled by the CO2 buffer system of
reactions (13) or (14). All wells in Kakkonda and
a low temperature well in Sumikawa plot be
tween curves (C) and (Q). Zeolitic minerals have
errors in the free energies (e.g., Hedenquist and
Browne, 1989), such that the positions of curves
(Q) and (C) cannot be well-defined at present.
The Pco2 of wells in the Kakkonda and
Sumikawa systems may be buffered by reaction
(10) or (12). The Pco2 of Japanese geothermal
fluids which are not greatly affected by boiling or
mixing are probably controlled by silicatemineral buffers described by reactions
through (14).(10)
Chemical condition of the neutral pH geother
mal reservoir
The pH of the neutral pH, NaCl-dominant
geothermal fluids can be approximated by reac
tion (3) (Fig. 8). Activity ratios of major cations,
such as (Na+) / (K+), (Na+) / (Ca2+) and
(K+)2/(Mg2+), are likely controlled by a silicate
mineral assemblage above 150°C or more conser
vatively between 200° and 300°C (Fig. 9).
Japanese geothermal fluids are also saturated
with respect to anhydrite and calcite (Fig. 11). At
tainment of these equilibria suggest that the
chemical compositions of geothermal fluids are
roughly predictable if two parameters are provid
ed, as stated by Arnorsson et al. (1983a). For ex
ample, if temperature and Z (Na + K) concentra
tion are provided, pH and activities of Na+, K+,
Ca2+, Mgt+, SO4 and C02 3 ions can be
calculated directly from Figs. 8, 9 and 11, since
Na+ and K+ ions are the most abundant among
the Na and K-bearing species. Si02 (aq) concen
tration can be fixed by the saturation with
respect to quartz or chalcedony. The concentra
tion of the most abundant anion, Cl-, can be
also calculated from a charge balance equation
after iterative calculations involving complex
aqueous species.
The oxygen fugacity range of fluid in
equilibrium with volcanic rocks between 200°
and 500°C were experimentally determined by
Kishima (1989) as a function of temperature.
The oxygen fugacities of neutral pH geothermal
fluids almost agree with the experimental data
(Chiba, unpublished data). This means that the
oxygen fugacity as well as the full chemical condi
tion in the neutral pH geothermal reservoir can
be roughly predicted if two parameters are pro
vided. The chemical condition predicted in this
way can be used as a constraint for chemical
modeling of geothermal systems, as a chemical
condition for simulation calculation of transport
of a particular element, and so on.
Solution and gas equilibrium 351
CONCLUSIONS
The aqueous species and pH of fluid discharg
ed from Japanese geothermal systems are approx
imately in equilibrium with their reservoir
minerals. The reservoir rock type as well as the
salinity of fluid does not affect the equilibrium
composition of aqueous species. Boiling with as
much as 10% steam loss in the geothermal reser
voir only slightly affects the activity ratios of
aqueous species. The pH and (K+)/(H+) activity
ratio of fluid in the geothermal reservoir is likely
controlled by K-bearing silicate minerals.
(Na+) / (K+) and (Na+) / ( activity ratios
are thermodynamically approximated by reac
tions between albite and K-feldspar, and be
tween albite and anorthite (or Ca-zeolites), re
spectively. The (Mg2+)/(K+)2 activity ratio of
Japanese high temperature geothermal fluid can
be represented by reaction among Mg-chlorite,
K-bearing silicate minerals and quartz, though at
lower temperatures other reactions may be
responsible for controlling this ratio. In addition
to silicate minerals, Japanese geothermal fluids
are saturated with respect to anhydrite and
calcite.
The partial pressure of CO2 is controlled by
reactions involving calcite, K-bearing silicate
minerals, and albite or Ca-zeolite in geothermal
systems not affected by steam loss and dilution.
Equilibrium between CH4, C02 and H2 is attain
ed at high temperatures but not maintained to
lower temperatures in most Japanese geothermal
systems, due to steam and gas loss, and
sometimes dilution. The H2/H2S ratios of some
fluids are likely equilibrated with Fe-bearing
minerals, though detailed studies of Fe-bearing
minerals are required to confirm the buffer
system of sulfur-bearing gases. Gas composi
tions are very good indicators of processes in the
geothermal reservoir, such as boiling and dilu
tion.
The characteristics of individual geothermal
systems can be described by the degree of partial
equilibrium attained in each system. Such
descriptions can become more quantitative if
kinetic data of reactions can be incorporated. Finally, the major aqueous composition and
pH of Na-Cl type geothermal fluids in Japan are
predictable if two variables (e.g. temperature
and one of the activities of the major com
ponents) are provided. Also, the results of this
study provide the basic knowledge necessary to
investigate the processes governing hydrother
mal mass transfer in the shallow part of the
earth's crust.
Acknowledgments-I wish to thank W. F. Giggen
bach, J. W. Hedenquist, N. Takeno and H. Shigeno
for their constructive comments on an earlier version
of this manuscript. I also thank K. Shimada, T.
Yokoyama, A. Ueda, T. Takenaka, Y. Yoshida,
Kyushu Electric Co., Mitsubishi Material Co. and
Idemitsu Kosan Co. for supplying much of the un
published data that made this study possible.
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Solution and gas equilibrium
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|
Chiba (1990) - Attainment of solution and gas equilibrium in japanese geothermal systems.txt
|
Supplementary Information
Feasibility of 129I groundwater dating calibrated by both 81Kr and 4He for the assessment of deep geological repositories in Japan
Tomoko Ohta1,2,*, Takuma Hasegawa1, Wei Jiang3,4, Guo Min Yang3,4, Zheng -Tian Lu3,4, Yasunori Mahara5
1. Civil Engineering Research Laboratory (Sustainable System Research Laboratory), Central Research Institute of Electric Power Industry, Abiko , Chiba, 270 -1176, Japan
2. Department of Nuclear Technology, Nagaoka University of Technology, Kamitomioka , Nagaoka, Niigata, 940 -2188, Japan
3. CAS Center for Excellence in Quantum Information and Quantum Physics, School of Physical Sciences, University of Science and Technology of China, Hefei, 230026, China
4. Hefei National Laboratory, University of Science and Technology of China, Hefei, 230088, China
5. Kyoto University, Sakyo -ku, Kyoto city, Kyoto 606 -8501, Japan
*Corresponding author: tomoohta@vos.nagaokaut.ac.jp
Figure captions
Table S1 Concentrations of tritium, 14C, TOC, δD, δ18O, and trace elements
Figure S1 Sedimentary age and geohistorical map of the sampling site.
*reorganization quoted from Ikawa et al.39
Figure S2 Concentrations of major ions in the groundwater
(a) Hexa -diagram, (b) Piper diagram
Figure S3 Schematic drawings of groundwater sampling at the sampling site
(a) Ground sampling, (b) overview of in situ sampling, (c) detailed in situ sampling
Figure S4 Correlation between iodine and TOC in the groundwater.
■: 127I,●: 129I
Figure S5 Correlation between I/Br and 1/Cl.
■: groundwater, ●: seawater, pink area: biogenic ratio of I/Br ( ≥0.04±0.02)
Figure S6 Correlation between iodine and boron.
(a) concentration of 127I and B, (b) concentration of 129I and B , T: only two points have low 129I concentrations, A:
regression curve excluding two points of T, (c) 129I/127I ratio and B, (d) 129I/127I/B ratio and B, ■: groundwater, ●:
surface seawater estimated from 129I/127I value (150 ×10-14)15and concentration of B in surface seawater (4.5 mg
L-1)
Table S1*
Geology Depth3H*,** TOC TC TICδ13C14C* δD*δ18O* Br* I/Br Cl* 1/Cl B
m TU mg L-1mg L-1mg L-1pMC ‰‰ mg L-1mol mol -1mg L-1mol -1mg L-1
Upper Sarabetsu F. 90.7 -99.7 DL 12 100 90 14±0.1 -81 -11 1.7 3.6 150 233 0.5
90.7 -99.7 DL 120 100 11±0.1 -82 -12 1.8 160 219
90.7 -99.7 DL 100 91 -1.6 9±0.1 -80 -12 1.7 160 219
214-215 DL 15 110 92 -0.5 2±0.0 -80 -12 5.8 2.0 570 61 0.9
Lower Sarabetsu F. 306-307 DL 10.2 61 47 22±0.1 -69 -11 1.3 130 269
306-307 DL 8.0 46 36 -13.8 15±0.1 -68 -10 2.9 0.28 280 125 0.5
337-348 DL 107 42 34 -8.9 15±0.1 -69 -11 1.6 0.34 150 233 0.3
Upper Yuchi F. 476-477 DL 118 170 130 +0.8 12±0.1 -70 -10 18 1.7 1600 22 3.8
613-614 DL 319 480 420 3±0.1 -46 -6 98 0.42 8500 4 3.8
613-614 DL 540 420 +2.7 ＜0.4 -45 -6 99 9300 4
715-716 DL 281 1100 730 +1.1 1±0.0 -13 1 180 0.29 17000 2 11
Lower Yuchi F. 943-944 DL 226 900 710 6±0.0 -11 1 190 0.22 17000 2 11.8
943-944 DL 900 620 -0.2 1±0.1 -11 1 190 17000 2
943-944 DL 870 640 1±0.0 -10 2 190 17000 2
943-944 DL 790 530 -0.3 ＜0.4 -10 2 180 17000 2
1143 -1144 DL 295 980 710 3±0.1 -14 2 170 0.31 16000 2 13
Sea water*** 0.67-0.84**** 67 0.00055 19350 1.8 4.5
* Hasegawa et al.25** DL: 0.1 TU *** Br, Cl in seawater: Mahara et al.1, I : Elderfield and Truesdale30
**** Halewood et al.39
*Hasegawa et al.24, **DL: 0.1 TU, ***Br, Cl in seawater: Mahara et al.1, I: Elderfield and Truesdale29, **** Halewood et al.38
Fig. S1Yuchi F.Late Pliocene
Approximately
2 Ma
Sea
Yuchi F.Early Pleistocene
Approximately
1 Ma
Bay to
lagoon
Sarabetsu F.
Yuchi F.Middle Pleistocene
Approximately
0.8~0.15 Ma
Lagoon to river
Yuchi F.0.018 Ma ~ Recent
Alluvium
Sarabetsu F. Sarabetsu F.Meteoric waterSampling year
(2017 -2018)
Period of last glacial
Mixing between meteoric
water and fossil seawater
fossil seawatermeteoric water since
last glacialAlluvium
*reorganization quoted from Ikawa et al.39
Fig. S2(a) (b)715-716m
943-944m
1143 -1144m90.7-99.7m
214-215m
306-307m
337-348m
476-477m
613-614mYuchi
Formation Sarabetsu
Formation
upper packer main valve
strainer
lower packer storage pumplifting pipe
perforation
interval radio -Kr
extraction
system
groundwater
sampling
radio -nuclides (129I,
3H, 14C), major
elements, minor
elements, stable
isotopesreducer
(a)In-situ groundwater
sampling
stable noble gases: for
4He datingTank for
ethanol
compressorElectric
winchtube
(b)Piston sampler
Check valveCupper tube
Fig. S3
(c)Piston sampler
Cupper tube (with
pure water)Piston
valve
valve
Check valve
Check valve closeethanol
①The process of lowering sampler with a winch pressurized
Pistondecompression
Check valve open
②In-situ samplingPiston
Check valve close
③ after in -situ samplinggroundwater
Cupper tube Cupper tube
(with in -situ
water)pressurized
Fig. S3ethanol
Fig. S4
Fig. S5
Fig. S6
y = 0.1218x + 10.199
R² = 0.0278y=22.53 ln(x)+23.81
R2=0.9561
TA
y = 5.241x+2.478
R² = 0.9593
y = 10.11x-0.995
R² = 0.9393
|
Ohta I129 sup41598_2024_66250_MOESM1_ESM.txt
|
Geochemical Journal, Vol. 7, pp. 123 to 151, 1973
Hydrogen and oxygen isotopic ratios and major element
chemistry of Japanese thermal water systems
OSAMU MATSUBAYA', HITOSHI SAKAI', ISAO KUSACHI2 and
HIROSHI SATAKE3
Institute for Thermal Spring Research, Okayama University,
Misasa, Tottori-Ken 682-02', School of Education, Oka
yama University, Tsushima, Okayama, Okayama-Ken 7002,
Department of Chemistry, Tokyo Kyoiku University,
Otsuka, Bunkyo-ku, Tokyo 1123, Japan
(Received August 13, 1973)
Abstract-More than 140 thermal waters of Japan were studied isotopically and chemi
cally. Highly saline brines at Arima and Takarazuka, Hyogo-Ken, and Ishibotoke near
Osaka indicate wide ranges of 5180 and 51) values from meteoric values of 5`0 = -8.2 and
SD = -50.0 %o SMOW to highly shifted values of +6.5 and -27.8 %o, respectively. The
isotopic values of these brines vary proportionally with chloride concentration irrespec
tive of temperature, carbonate concentration or locality. These saline waters are iso
topically and chemically best explained as the mixtures of local meteoric waters and a
saline brine of C 1 = 43,700 ppm, 5180 = +8 %o and 8D = -30 to -25 %o. The latter is most
likely the "residual magmatic, metamorphic or geothermal" fluid associated with upper
Cretaceous rhyolitic and granitic rocks and Ryoke metamorphic rocks in which these
brines are found. Thermal waters at Ikeda and in adjacent areas, Shimane-Ken and at
Senami, Niigata-Ken, are similar to the Arima brines in the isotopic and major element
chemistry, but are much more diluted by the respective local meteoric waters.
Many of the thermal waters along the ocean coasts are isotopically intermediate
between oceanic and local meteoric waters and are considered to be mixtures of the two
types of water. As a result of hydrothermal mineral-sea water interaction, the coastal
thermal waters differ considerably in the chemistry from fresh sea water and are typical
of Na-Ca-Cl type. The coastal thermal waters isotopically and chemically may be similar,
if not the same, to submarine hydrothermal ore fluids responsible for the Kuroko type
mineralization. The isotopic values and their relationship to salinity, however, widely
differ from one system to another depending on the hydrogeological conditions of
each system. The coastal thermal waters at Ibusuki of Ata Caldera, Kagoshima-Ken,
for instance, are significantly affected by the waters from three crater lakes, Lake
Ikeda, Unagi-Ike and Kagami-Ike, in which the 5180 and 6D values are meteorologically
balanced at such high values as -2.6 and -19.4%o, respectively.
Acid to neutral thermal waters of volcanic affiliation indicate varying degrees of
isotopic shifts, but they are supposed from their 6D values to be essentially derived from
recycled meteoric water.
123
124 O. MATSUBAYA et al.
Many thermal waters of neutral chloride type in the "green tuff" regions of the
inner Honshu also are simple meteoric in origin without showing any significant isotopic
shifts, although the waters are relatively high in salinity and SO4/Cl ratios.
INTRODUCTION
Most of the thermal spring systems in Japan are closely associated with Tertiary
to Recent volcanic and plutonic rocks. Most of the springs are in the "green tuff
regions that contain Miocene geosynclinal sediments with abundant submarine pyro
clastic rocks and lava flows (Fig. 1). The "green tuffs" also are the host rocks for many
"Kuroko" ore deposits. The thermal waters range from acid sulfate-dominant waters
of volcanic association to neutral chloride or sulfate-dominant thermal waters of
varying chemical compositions and concentrations.
0
a
o
35 36-41P 231 7-14
Nobo Penn.!
18-22 "`
o06
42-44 '105-17 ° 08,..9 0 4
88 -7104
.0 111 • 61-s7
140J41 ': 73-80 Kii Penin.o
O
112-125
-126-132
133-139ri:.
68 `
a5-67
i:u Penin.3-5
Fig.1. Distribution of thermal water systems of Japan.
• : Thermal water systems of the present study.
o : Other major thermal water systems.
Solid and dotted lines define the outermost margin of the "green tuff
regions, the dark areas are underlain by typical "green tuff" formations.
Chemical and geological evidence led many Japanese investigators (e.g. NAKAMURA,
1962; UZUMASA, 1965; YUHARA and SENO, 1969) to believe that the waters and
dissolved chemical species of Japanese thermal water systems originate from at
least three major sources, that is, volanic, oceanic (present-day), and fossil or
connate, the last mentioned being most likely in the "green tuff" regions. However,
D/H, 180/160 and major element chemistry of thermal waters 125
chemical criteria have not been powerful enough to discriminate them from each
other.
Conversely, oxygen and hydrogen isotopic studies are useful in the genetic clas
sification of thermal waters (CRAIG, 1963, 1966; CLAYTON et al., 1966; HALL and
FRIEDMAN, 1963; RYE, 1966; OHMOTO and RYE, 1970; MIZUTANI and HAMASUNA, 1972).
Deuterium concentrations of Japanese thermal water systems have been extensively
measured by KOBAYAKAWA (1960), but until now the simultaneous analyses of oxygen
and hydrogen isotopic ratios are not abundant.
We have analyzed the isotopic ratios of water and sulfate in more than 140 ther
mal waters of various types. The WC values of carbonate in some saline brines at
Arima and in adjacent areas were also analyzed. Major cations, anions, and some trace
element concentrations of most of these waters were analyzed by KUSACHI (in prepara
tion) and the results were partly discussed by HENMI and KUSACHI (1966). The
present paper reports the oxygen and-hydrogen isotopic results and major element
chemistry (Na+, K+, Cal, Mg", SO4-, C1) of the waters. Isotopic ratios of the ther
mal waters at Arima and in related areas, and at ocean coasts will be discussed in some
detail with special reference to Na-Cl type magmatic hydrothermal fluids and to Na
Ca-Cl type ore solution that presumably are responsible for the Kuroko type minerali
zation. The isotopic geochemistry of sulfate in thermal waters will be discussed in an
accompanying paper (SAKAI and MATSUBAYA, in preparation).
EXPERIMENTAL AND RESULTS
Oxygen isotopic ratios of water were measured by the conventional C02-H20
exchange technique (EPSTEIN and MAYEDA, 1953). Deuterium determinations were
carried out on H2 obtained by passing water vapor over heated uranium metal
(FRIEDMAN and SMITH, 1958). Isotopic analyses of oxygen and hydrogen were done in
the Institute for Thermal Spring Research, Okayama University, while some of deu
terium analyses were carried out in the Department of Chemistry of Tokyo Kyoiku
University. Carbonate in the saline brines at Arima and in adjacent areas were precipi
tated as BaCO3 at the sites of collection. The precipitates were brought back to the
laboratory in sealed bottles and subjected to the carbon isotopic analyses by the
conventional H3PO4 technique. The isotopic ratios are reported in 5X values:
8X = (Rsample _ 1) X 1,000 o/00 R
standard
where X = 180, 13C or D, R = 180/160, 13C/12C or D/H. The standard for oxygen and
hydrogen is SMOW standard from the International Atomic Energy Agency and that
for carbon is PDB. The analytical errors are ±0.1 %o for oxygen and carbon and
±0.7%o for hydrogen.
126 0. MATSUBAYA et al.
The isotopic values of hydrogen and oxygen (and sulfur) are listed in Appendix 1
together with the chemical compositions of the major elements, pH, and temperature
of the waters. Localities of the thermal systems are shown in Fig. 1 by closed circles
and numbers. Brief description of some of these thermal.. systems will be found in the
following sections. For further information, the readers may be referred to UZUMASA
(1965), NAKAMURA (1962), YUHARA and SENO (1969) and the references therein. The
isotopic ratios of meteoric waters at or near the thermal springs are shown in Appendix
2. The oxygen and hydrogen isotopic results are plotted in Figs.2a, 2b and 5.
-1U
I I 1 I I-5I I I Z.+5
so=88180+10
• 0
00•
j
e°U eI
-50 0
LocalityI busukiN
arikawa
KagoshimaSouthern Eastern
Kii Izu
peninsula peninsulaNoto
peninsulaAnima
TakarazukaIshdlotoke
wearI 4 e • a
Meteoricwater ® 6 A m m
Fig.2a. Isotopic composition of thermal and meteoric waters at Arima, Ta
karazuka, Ishibotoke, and ocean coasts. The plot for Ibusuki includes five
points not listed in Appendix 1.
-1510S~bSMOW°~aa
--5
SD=66180+10
J
O
OF f--50
sLocalities BeppuareaManza Cent
KyusrhaulTar~i~ a jacentHakoneGreen tuff
regions o
inner Honshu
water l • 4 s
Meteoric water
Fig.2b. Isotopic composition of thermal and meteoric waters in Quaternary
volcanic areas, in the "green tuff" regions of inner Honshu and in other areas.
D/H, 180/160 and major element chemistry of thermal waters 127
DISCUSSION
Brines and carbonated waters at Arima, Takarazuka and Ishibotoke Arima (89
104)* is on the northern slope of the Rokko Mountains north of Kobe. The Rokko
Mountains are composed of upper Cretaceous granitic rocks, which intruded into Ari
ma rhyolite of a similar age (KASAMA, 1968). The Arima thermal water system is in
the Arima rhyolite close to the major thrust-faults of this area (NAKAMURA, 1962).
Three types of water are noted in this area (TSURUMAKI, 1965:) (1) saline brines
of high temperature (ca. 97'C), (2) saline brines of low to medium temperature (18
60°C) but of high carbonate (HC03 + C02) concentrations and (3) CO2 -rich waters
of low salinity and temperature. More than twice as high a salinity as that of average
sea water has been reported for one of the first type water (IKEDA, 1949).
Takarazuka (88), Hyogo-Ken, is on the eastern extention of the Rokko granite
and is characterized by C02-rich mineral waters of high salinity (TSURUMAKI, 1965)
and is similar to the second type of water from Arima.
Saline carbonated waters at Ishibotoke (81 87), Kawachi-nagano south of Osa
ka are from the commercial wells for the C02-production which were drilled into
Ryoke metamorphic rocks of Cretaceous age underlying Quaternary sandy deposits
(Osaka formations). The CO,-rich brines have been found only in the Ryoke met
amorphic rocks (TSURUMAKI, private communication). These brines also are similar to
the second type at Arima.
MIYAKE et al. (1955) reported that excellent linearity exists between Cl and
other ions such as I-, Br-, HBO, and Ca" in some high temperature brines at Arima.
However, TSURUMAKI (1965) demonstrated that the linear relationship between Cl
and other ions except for carbonate holds more extensively among all the three types
of water including those at Takarazuka and Ishibotoke, while carbonate concentration
apparently is controlled by the temperature of water.
As is seen in Fig.2a and Appendix 1, these waters show wide ranges of oxygen
and hydrogen isotopic ratios; the C02 rich waters of low salinity at Arima (93, 103,
104) and Ishibotoke (87) are isotopically almost identical with the local meteoric
waters (M32--36), while the saline waters exhibit varying degrees of isotopic shifts.
Figure 3 indicates that the oxygen isotopic shift is proportional to the chloride con
centration irrespective of the locality, temperature and total carbonate concentration.
Though not shown, the hydrogen isotopic shift also is found to be almost linearly
related to Cl concentration. As is shown in Table 1, the S13C values of total carbonate
(CO2 + HC03) in these brines are in a relatively narrow range from -2.9 to -8. 1 %o
PDB. The carbonates in the dilute C02-rich waters at Arima and Ishibotoke, however,
are lighter than those in the saline brines from these and Takarazuka areas.
The present isotopic results as well as the chemical evidence by the previous
investigators strongly support the view that the brines at Arima are admixtures of a
deep saline brine and local meteoric waters (TSURUMAKI, 1965). The brines at Taka
* Numbers in parentheses are sample numbers in Appendixes 1 and 2.
128 0. MATSUBAYA et al.
+5
0
-56180
0
F' 00 •
7O0
•'O
O11
OZ'Z11
0 Arima, high temperature brines
• : Arima, low to medium temperature brines
and meteoric waters(at Cr=O)
0: Arima, the brine of the maxirrwm sainity(see the text)
0 : Takarazuka brine
0: Ishibotoke brines and meteoric waters(Cl-=0)
molI -10 CIr m O
Fig.3.0 500 1000
Relationship between 5180 and chloride concentration in brines and
carbonate waters at Arima, Takrazuka and Ishibotoke.
Table 1.
the brinesCarbon isotopic ratios of total carbonate (CO2 + HCO3) in
and dilute carbonate waters of Arima, Takarazuka and
Ishibotoke
No. 1) Locality tocEC022)
m mol/1 07
m mol/1S13C PDB
00
81
85
87
88
94
95
96
97
98
99
100
101
102
103
104Ishibotoke-1
Ishibotoke-5
Ishibotoke-7
Takarazuka
Arima-6
Arima-7
Arima-8
Arima-9
Arima-10
Arima-11
Arima-12
Arima-13
Arima-14
Arima-15
Arima-1621
21
20
20
b.p.3)
b.p.
b.p.
b.p.
61
51
48
30
21
18.5
15.286
59
49
76
0.4
2.0
3.4
2.2
9.2
16
10
4.3
9.8
31
5.6417
26.4
16.6
350
1,061
803
634
540
563
41.8
928
160
408
1.3
0.6-4.8
-7.3
-8.0
-3.2
-4.1
-5.6
-6.2
-4.2
-5.4
-6.8
-5.7
-3.5
-2.9
-7.7
-8.1
1) Numbers in Appendix 1.
2) Total carbonate concentration, not
values.
3) The waters are near boiling point.reliable better than±10% of the
D/H, '$0/160 and major element chemistry of thermal waters 129
razuka and Ishibotoke can be explained by essentially the same model. The deep
brines in the three areas seem to have the same isotopic and chemical characteristics.
Cationic compositions of these waters vary considerably within the Arima district as
well as between Arima and other districts. However, an inspection of Appendix 1
would reveal that K/Na and Ca/Mg ratios are generally higher in high temperature
brines at Arima than in other low temperature brines. The cationic compositions of
these brines may largely be controlled by the temperature at which the brines were
equilibrated with wallrock minerals (e.g. ORVILLE, 1963).
The deep saline brine may have the 5180 value of +8%o (shown by the crossed
open circle in Fig.3), which corresponds to the water of the maximum salinity (Cl =
43,700 ppm) so far reported (IKEDA, 1949). Using Fig.2a, the 5D value of the brine
may be estimated to be from -30 to -25 %o, depending on whether the two points
at the far top-right of the plot are taken into account or not.
The high isotopic values and salinity of the brines and their close association with
the upper Cretaceous granitic or metamorphic rocks suggest that the brine reservoirs
might have been derived from the residual "magmatic" or "metamorphic"* waters
entrapped in these rocks (NAKAMURA, 1962). The 5180 and 5D values of the "magmatic
fluid" would largely be determined by the isotopic equilibria between the fluid and
magma. The 5180 values of quartz in Hiroshima granite batholith, in a granite of the
Ryoke metamorphic zone near Hiroshima and in the granitic complex at Ibaragi near
Osaka range from +9 to +12%o (MATSUHISA et al., 1972; MATSUHISA et al., 1973;
H. HONMA, personal communication). The 5180 values of water in isotopic equi
librium with these quartzes at 600°C would be from +7 to +11%o according to
the latest quartz-water isotopic temperature scales (SHIRO and SAKAI, 1972; CLAYTON
et al., 1972). TAYLOR (1967) also suggested that "magmatic water" would have 5180
values from +7.5 to +9%o and that metamorphic water would have higher values.
The inferred 5180 value of the deep saline brine at Arima and in adjacent areas is well
within the possible range of "magmatic" and "metamorphic" waters.
The range of SD values of "magmatic and metamorphic fluids", on the other
hand, is still ambiguous at present. It should depend on the ultimate source of water
(juvenile, meteoric, oceanic or mixture of these), relative amount of hydrogen parti
tioned among coexisting phases, and the temperature of crystallization and, thus, it
could vary widely. The SD value, -30 to -25 %o of the present saline brine may be
taken as an example of SD value of "magmatic and metamorphic waters".
Alternatively, local meteoric waters might have come to attain such high isotopic
values and salinity after being subjected to interaction with surrounding rocks at
elevated temperatures, and/or to kinetic evaporation condensation (CRAIG, 1963).
However, it should be noted that the studied areas, especially Takarazuka and Ishi
* Fluid associated with rocks during metamorphism (WHITE, 1957b).
130 0. MATSUBAYA et al.
botoke, are presently not active geothermal areas. The high temperature brines at
Arima are found only within a narrow area of 300m across. Furthermore, the isotopic
shifts are parallel with salinity but not with temperature of brine. This is contrasting
to the Salton Sea geothermal brines (CRAIG, 1966) and thermal waters of the present
study from Quaternary volcanic areas in which the oxygen isotopic ratios are not
linear function of Cl or salinity as will be described later. These features suggest that
the brines are of "fossil" geothermal origin.
Although the choice between the two models described above is ambiguous at
present, the "magmatic" and "geothermal" brines may be essentially the same, if the
"magmatic" fluid is ultimately of meteoric origin.
The carbon dioxide in the brines seems to be inherent in the deep saline brine
(TSURUMAKI, 1965). The high temperature brines at Arima are depleted in carbonate,
probably because carbon dioxide was boiled off from the heated brines and calcite
was precipitated. The higher S13C values in the low temperature saline brines of
Jadani (102), Gekko-en (101) and Takarazuka (88) may closely represent the original
carbon isotopic ratios of carbonate, while the lighter isotopic values in the C02-rich
shallow waters may reflect some contamination by organic carbon, or isotopic frac
tionation in the separation of gaseous carbon dioxide from solution. The higher
S13C values in the three brines mentioned above are rather similar to those of some
marine limestones and hydrothermal carbonates (M. WATANABE, personal communica
tion, 1972), but are quite unlike organic carbon, of which VT values are generally
lower than -25%o.
Coastal thermal water systems Many of the thermal water systems of Japan
which lie close to and along the ocean coasts exhibit characteristic features in the
chemistry and isotopic ratios as the results of the mixing of sea water into the thermal
systems. Well studied systems, where "partial sea water origin" of the thermal waters
was suggested on the basis of chemical evidence, are Ibusuki (133 138) and Narikawa
(139) of Kagoshima-Ken (HATAE et al., 1965), Shirahama (78, 79) and Katsuura (73,
74), southern Kii Peninsula of Wakayama-Ken (NAKAMURA, 1962) and Atami (59, 60)
and Ito (62), Izu Peninsula of Shizuoka-Ken (NAKAMURA et al., 1969; MUROZUMI,
1970). "Partial sea water origin" of the Ibusuki thermal waters was also suggested
by oxygen isotopic ratios (TARUTANI et al., 1972). More recently, a detailed study of
the isotopic ratios of water and sulfates and the chemical composition of the thermal
waters at Shimogamo on southwestern coast of Izu Peninsula led MIZUTANI and
HAMASUNA (1972) to conclude that those thermal waters are admixture of local me
teoric water and deep-seated hydrothermal brines of sea water origin.
Thermal waters from the above mentioned coastal hot-springs except for Shimo
gamo were isotopically analyzed in the present study. Figure 2a indicates that the
isotopic values of these coastal thermal waters are intermediate between those of
D/H, '80/'60 andmajor element chemistry of thermal waters 131
SMOW and respective local meteoric waters. This strongly supports the view that sea
water is involved in the hydrothermal systems. Inspection of Appendix 1 shows that
these waters are typically high in salinity and of Na-Ca-Cl type chemistry. The SO4/Cl
ratios in them are lower than in fresh sea water but the 5345 values of sulfate are
practically the same. Thermal waters at Kaike (36) of Tottori-Ken and Kagoshima
(132) of Kagoshima-Ken can also be regarded as diluted sea water. A high Li/Na
ratio in Kaike (3.6) also led HENMi and KUSACHI (1966) to the same conclusion.
Figure 4 is a 5180 vs. Cl concentration plot of the coastal thermal waters of
the present study. A similar plot would be obtained for SD. The 5180 values of the
respective local meteoric waters are plotted at Cl = 0. Figure 4 indicates that these
thermal waters rarely are simple mixtures of fresh sea water and local meteoric waters
but either of them or both must have been altered isotopically before and/or after
the mixing. Some examples will be discussed below.
0
-1
-x
-38
_5j
Qlee
8•----Ao-`-6
cs
O : lbusuki • : Atami and Ito
6 : Atagavva m : Shirahama and Katsuura
0 : Wakura O : Sea water
o : heavy meteoric water at Ibusuki
: normal meteoric water at lbusuki(see the text)
A : sea water of oxygen isotopic shift
A : meteoric waters of relatively high salinity
Sold and dotted lines: mixing lines of sea water
and meteoric water(see the text)
100 200 300 400 500
C1'-concentration m mol.
Fig.4. Relationship between 5180 and chloride concentration in some coastal
thermal waters.
Points on the ordinate represent the respective average meteoric waters (for
Ibusuki, see the text and explanation in the figure). Solid and dashed lines;
mixing lines of sea water and meteoric water (see the text).
Ibusuki and Narikawa are along the coast of Kagoshima Bay and are in the
large caldera of Ata Volcano (about 20km in diameter), of which the eastern half
underlies Kagoshima Bay (ARAMAKi and Ui, 1966). The thermal waters are related to
Quaternary volcanism, which formed a series of craters and cones within the old
Ata Caldera (TSUYUKI, 1965). The thermal waters emerge out or are found by
132 0. MATSUBAYA et al.
drilling along the coastal line which partly is composed of the western rims of
several small craters. Two crater lakes give some complicated but interesting hydro
geologic features to the Ibusuki and Narikawa thermal water systems (see below);
Lake Ikeda, 3 4km in diameter and about 250m of water depth, is near the western
rim of the old Ata Caldera and 7 8 km west of K,agoshima Bay, while the other,
Unagi-Ike (Lake Eel), is about 1 km across and about 1.5 km southeast of Lake Ikeda.
As is seen in Fig-4, the 5180 vs. Cl plot for these thermal waters does not fit a
mixing line between the average sea water and the average of the three meteoric
waters (M-38, -39, -41) (shown by a solid square at Cl = 0 in Fig.4), which were
collected from the shallow wells between the coast and Unagi-Ike and isotopically
should represent the precipitation within this area. Instead, the isotopic values of the
Ibusuki and Narikawa thermal waters fall within a group of mixing lines (A and A',
in Fig.4) between the average sea water and meteoric waters of which 5180 values
should be from -5 to -6%o. Therefore, we looked for surface waters of such heavy
isotopic ratios and found that waters from Lake Ikeda, Unagi-Ike, Kagami-Ike (a small
maar), and a well on the coast of Lake Ikeda are isotopically quite heavy (shown by
open squares at Cl = 0 in Fig.4).
In Fig.2a the 5180 vs. 6D plot of these waters from the lakes and wells forms a
straight line of a slope (6D/6180) of 5.5. The slope is quite similar to that experi
mentally obtained by CRAIG et al. (1963) in the kinetic evaporation of limited amounts
of water at normal temperatures. Thus, the isotopic values of Lake Ikeda and its
surrounding high land area are controlled by the isotopic and mass balance between
the rapid kinetic evaporation, drainage, and precipitation. It is likely that the heavy
waters in Lake Ikeda and Unagi-Ike significantly contribute to the aquifers of the
coastal thermal waters. Rather erratic scatter of the 5180 vs. 6D plot in Fig.2a for
the Ibusuki and Narikawa thermal waters may be due to variable mixing ratios of the
oceanic and two types of meteoric waters.
Alternatively, as a result of the oxygen isotopic exchange with wall rock minerals,
the hydrothermal brines at Ibusuki may have a 6180 value higher than sea water. If
this is the case the mixing model represented by the dotted line (B) in Fig.4 may
resolve at least some of the erratic isotopic pattern at Ibusuki.
Figure 4 indicates that the isotopic trend at Shirahama (78, 79) is rather similar
to that observed at Ibusuki (133 138) and may be similarly explained by the mixing
model represented by B. On the other hand, the thermal waters at Atami (59, 60)
and Ito (62) of Izu Peninsula lie below the respective mixing line between the average
sea water and the local meteoric waters in Fig.4. These thermal waters may be mix
tures of sea water-derived brines of normal isotopic values and meteoric thermal
waters of relatively high salt concentration (solid line C in Fig.4). Thermal waters
numbered 63 to 67 have sulfate isotopically quite different from sea water sulfates
(Appendix 1) and may represent such meteoric waters of high salt concentration.
D/H, 180/160 and major element chemistry of thermal waters 133
Close inspection of the major element chemistry of these coastal thermal waters
(Appendix 1) reveals that these waters are not simple diluted sea water but are
generally depleted, in terms of chloride-normalized concentration, in Na+, Mg" and
SO' and enriched in Ca" and K+ relative to sea water. Apparently this is due to
chemical interaction between sea water and wallrocks in the thermal water systems.
Albitization and Mg-chloritization of calcic minerals would considerably increase the
Ca" concentration in the solution with expenses of Na+ and Mg". As a result, pre
cipitation of anhydrite would be enhanced, leading to a depletion of sulfate in the
solution without any significant change in the isotopic values of the dissolved sulfate
(see MIZUTANI and HAMASUNA, 1972; SAKAI and MATSUBAYA, in preparation). The
isotopic and chemical characteristics of the coastal thermal water systems discussed
above may be similar, if not the same, to those of the hydrothermal systems responsi
ble for the Kuroko ore deposits (SAKAI et al., 1970).
Wakura (22) is quite similar in its high salinity and major element chemistry to
other coastal thermal waters of "partial sea water origin". However, as is seen from
Fig.2a, the isotopic values of this water are shifted significantly along the CRAIG'S
meteoric plot and cannot be explained by any reasonable mixing model between sea
water and local meteoric water. The water of crystallization of gypsum in the Kuroko
type gypsum deposits about 50km north of Wakura has, on an average, the peculiar
isotopic values of 5180 = -1.1 and SD = -32.8 %o, respectively (MATSUBAYA and SAKAI,
1973). It may be interesting to note that the thermal water at Wakura is isotopically
close to the water that should have been in isotopic equilibrium with the Noto gypsum
at 20 to 60°C (MATSUBAYA and SAKAI, 1973). However, implication of this fact is not
celar at present.
Volcanic thermal water systems Thermal water systems of strong volcanic af
filiation widely exist in the Japanese Islands, especially in the "green tuff" regions.
Tamagawa (1, 2), Akita-Ken, Kusatsu-Shirane (69 72), Gunma-Ken, Hakone (45
58), Kanagawa-Ken, Beppu (112 118) and others (119 123), Oita-Ken and Shi
mabara (140, 141), Nagasaki-Ken are typical of the volcanic thermal waters of the
present study. Ibusuki (133 138), Narikawa (139) and Kinko (132), Kagoshima
Ken, which also belong to this group, have already been discussed as the coastal
thermal water systems. Most of these waters are acid sulfate and acid chloride sulfate
type waters, although neutral chloride type waters also have been found in deep
aquifers of many volcanic and geothermal areas (e.g. UZUMASA and MUROZUMI, 1955).
In the present study, Hakone-5 and -8 (49, 52) are the examples of such deep
seated neutral chloride type waters (OKI and HIRANO, 1972).
The volcanic thermal waters of the present study exhibit varying degrees of iso
topic shift extending from the respective meteoric values and with a slope of 2 or
less. The isotopic pattern (Fig.2b) is essentially the same as that observed for the acid
134 0. MATSUBAYA et alL
thermal waters by CRAIG (1963) and indicates that the volcanic thermal waters all
are meteoric in origin. However, close inspection of the data in Appendix 1 reveals
that the dissolved sulfates in acid chloride sulfate waters are usually much higher
in S34S than those in acid sulfate waters (compare, for instance Beppu-1 5 (112
116) with Beppu-6 (117)). The S34S values of the latter are similar to the volcanic
sulfur and hydrogen sulfide of the respective volcanoes and the 5180 values to what
are expected for freshly formed supergene sulfate in ground waters (SAKAI and MATSU
BAYA in preparation).
Acid sulfate waters in volcanic areas are considered to be formed by the surface
oxidation of volcanic hydrogen sulfide which is fractionally separated from con
densible hydrogen chloride and water. The chloride condensate would become chlo
ride-rich acid waters with varying degrees of dilution by ground water (WHITE, 1957a;
ELLIS and MAHON, 1964). The present data are in general accord with the above
model. However, the isotopically heavy sulfates in acid chloride sulfate waters should
originate from other sources than volcanic hydrogen sulfide, probably from "fossil
sea water sulfates" in the "green tuff" formations which underlie most of the volcanic
areas of the present study (SAKAI, 1969; SAKAI and MATSUBAYA, in preparation).
NAKAI (1972) and ABIKO (1972) also reported that two types of sulfate which signi
ficantly differ in S34S value exist in some volcanic thermal waters.
Unlike the saline brines of Arima and other areas, the isotopic shifts of these
acid chloride sulfate waters are not linearly related to the chloride concentration (or
salinity). This would also imply that the dissolved salts in these waters cane from
more than a single source.
Thermal water systems of the Japan Sea side of Honshu Most of other thermal
water systems of the present study are from the "green tuff" regions of the inner
Honshu (Japan Sea side of Honshu), although some are located in older (Crataceous
to Paleogene) plutonic and volcanic rocks off the "green tuff" regions. The "green
tuff" regions of the inner Honshu are much less populated with Quaternary volcanic
rocks than the northeastern back-bone area, central Honshu and Kyushu. The "green
tuff" regions of Niigata-Ken are one of the oil and natural gas fields of Japan.
The thermal waters in the "green tuff" regions of the inner Honshu are generally
of neutral sodium chloride sulfate type chemistry with relatively high salinities and
S04/Cl ratios. Occassionally, sulfate exceeds chloride. The present data (Fig. 2b,
5 and Appendix 1) indicate that most of these waters are simply recycled meteoric
waters. It should be noted, however, that the sulfates in these waters are usually
high in the S34S values and are similar to the sulfates in the Kuroko ore deposits
abundantly emplaced in the "green tuff" formations (SAKAI et al., 1970; SAKAI and
MATSUBAYA, in preparation). The sulfates in the thermal waters are considered to be "fossil" sea water sulfates in the "green tuff" formations and contemporaneous to
D/H, 180/160 and major element chemistry of thermal waters 135
-12 -11 -10 6110%.
-9 -8 -7
16-I
F15 M-5 M-4
M-6 I-1 34
M-7
17 39 38
M-11 33 0 0;
14 22 12 / 10 /
11
13
/ SD=86180+10L1-6
40
o® -50
-60
-70d 8
D
80
Fig.5. 5180 vs. 5D diagram for thermal waters in Niigata-, Tottori and Shima
ne-Ken. o : Niigata-Ken; 0 : Tottori-Ken and Shimane-Ken. Numbers are
sample numbers in Appendixes 1 and 2.
the Kuroko sulfates (SAKAI, 1969; SAKAI and MATSUBAYA, in preparation).
The isotopic values of thermal waters collected within a small area may vary
along the CRAIG'S meteoric water plot, even if the waters are simple recycled meteoric
waters. Tottori 1 (26) and Tottori-2 (27) are only a kilometer apart and so are
Misasa 1 (33) and Misasa-2 (34). The isotopic variation observed within each district
should be ascribed to complicated features of local meteoric and hydrogeologic
conditions.
As is shown in the expanded scale of Fig.5, the high temperature waters at
Senami (7, 8, 9), Niigata-Ken, apparently show the isotopic shift and are different
from other waters (10, 14) located further inland. However, it is not clear at present
whether the original waters of Senami are isotopically similar to the Miomote river
(M-6) and, thus, to the inland thermal waters or to the two well waters (M-4, -5)
at Senami. Senami is presently not active "geothermal" area. The heat source may
be Tertiary rhyolite which occurs within five hundred meters below the surface. A
temperature of 120 to 130°C has been measured in a drill hole in the rhyolite (S. ABE,
written communication). Granitic rocks of upper Cretaceous age has been located
within a kilometer from Senami. The Senami thermal waters may be mixtures of the
brine of "fossil" geothermal or magmatic origin similar to Arima brines and recycled
meteoric water. Although the "green tuff" formations were not identified at Senami,
the 5345 values of sulfate in these waters strongly suggest that the recycled meteoric
waters acquired most of the dissolved salts from the "green tuff" formations.
Low to medium temperature thermal waters at Shigaku (38), Yugakai (39) and
136 0. MATSUBAYA et al.
Ikeda (40) are within a narrow district near Sanbe Volcano, Shimane-Ken. Yugakai
and Ikeda are related to granitic rocks of presumably Paleogene age, whereas, Shigaku
is in the Quaternary volcanic rocks of Sanbe Volcano. Their relation with the "green
tuff" formation is obscure. These low temperature waters are characterized by high
salinity and Na-Ca-Cl type chemistry and probably are similar in origin to the Arima
thermal waters (NAKAMURA, 1962). As is shown in Fig.5, the isotopic values of Yuga
kai and Ikeda exhibit a slight but definite isotopic shift from Tenjin River (M 11)
which may represent the local meteoric waters of Japan Sea side of southwestern
Honshu. Although the observed magnitude of the isotopic shift is only 10% or less
of that of the Arima brines, the salinity of these waters also is about 10% of that
of Arima. Therefore, these waters may be highly diluted versions of the Arima brines
as has been suggested by NAKAMURA (1962). Chemically Shigaku (38) is similar to
Yugakai (39) and Ikeda (40) and may be grouped together with them. The isotopic
shift in Shigaku is negligibly small owing to further dilution by meteoric water. The
slight isotopic shifts of these waters imply that either saline brines of high temperature
similar to those found at Arima may exist in the deeper part of the granitic and vol
canic complex of this area, or the reservoir of such brines has been almost exhausted.
CONCLUSIONS
Isotopically and chemically four types of thermal waters are characterized in
the present study. First and the most distinguished is the possible "fossil magmatic,
metamorphic or geothermal fluids" which are found in the Cretaceous granitic rocks
at Arima and Takarazuka, Hyogo-Ken and in the Ryoke metamorphic rocks at Ishi
botoke, south of Osaka. The highly saline waters are characterized by the high
isotopic values of water and high carbonate concentrations as well as by the trace
concentration of sulfate. The "fossil brine" at depth may have the isotopic values
of 5180 = +8 and bD = -30 to -25 %o and the chloride concentration of 43,700 ppm.
Highly diluted versions of this type of water also are found at Yugakai, Ikeda and
Shigaku of Shimane-Ken and at Senami of Niigata-Ken, where the thermal waters
are associated more closely with pre-Neogene granitic and/or Neogene volcanic rocks
than with recent volcanic activities.
Secondly, a number of thermal waters at the ocean coasts show intermediate
isotopic values between near SMOW and local meteoric waters as a result of the
invasion of sea water into the thermal systems. The first isotopic documentation of
such system was given by MIZUTANi and HAMASUNA (1972) at Shimogamo, southwestern
Izu Peninsula of Shizuoka-Ken. The present study confirmed some more examples
of this type of water at Ibusuki of Kagoshima-Ken, Shirahama, southern Kii-Peninsula
of Wakayama-Ken, Ito and Atami of eastern Izu Peninsula of Shizuoka-Ken and Kaike
of Tottori-Ken. Hydrothermal reactions between fresh or diluted sea water and wall
D/H, 180/160 and major element chemistry of thermal waters 137
rock minerals produce typical Na-Ca-Cl type waters of low Mg" and SO,' concentra
tions. The 634S values of SOT are similar to the fresh sea water sulfates, because
S04 is removed by anhydrite precipitation.
The third type is found in close affiliation with Quaternary volcanism. The
waters may be further classified into acid sulfate type and acid to neutral chloride
type volcanic waters. These waters indicate varying degrees of isotopic shifts, although
they are essentially simple local meteoric waters. However, sulfates in the acid to
neutral chloride sulfate type waters are isotopically much heavier than the supergene
sulfates in the acid sulfate waters. It is suggested that most of these heavy sulfates
are derived from the "green tuff" formations underlying most of these volcanic areas.
The fourth type of thermal water comprises the majority of the thermal waters
in the "green tuff" regions of the inner Honshu. These waters are simple local
meteoric waters but are characterized by relatively high salinities and S04/Cl ratios.
As will be discussed elsewhere (SAKAI and MATSUBAYA, in preparation), the sulfates are
isotopically heavy and similar to those of the Kuroko ore deposits in the "green tuff
regions. It is suggested that a considerable portion of the dissolved salts is supplied
into the recycled waters from the "green tuff" formations.
ACKNOWLEDGMENTS
We are greatly indebted to Drs. E. CHENEY of University of Washington, A. SASAKI of Geolo
gical Survey of Japan, Y. OKI of the Hot Spring Research Institute of Kanagawa-Prefecture, A.
YOSHITANI of Tottori University, K. KIGOSHI of Gakushuin University and H. HONMA of Okayama
University for their critical reading the manuscript and valuable suggestions. A part of the
deuterium determinations in this study were carried out in the Department of Chemistry, Tokyo
Kyoiku University under the supervision of Dr. S. MATSUO, now of Tokyo Engineering University
to whom our sincere thanks are due. Dr. M. TSURUMAKI of Osaka City University, Drs. T. TARUTANI
and A. KOGA of Kyushu University, Dr. Y. OKI, Dr. S. ABE of Tokyo Agricultural and Engineering
University, Drs. M. SAKANOUE and H. SAKAMOTO of Kanazawa University, Dr. M. KAMADA
of Kagoshima University, Mr. J. HIRABAYASHI of Tokyo Engineering University and Dr. M.
ICHIKUNI of Tohoku University helped us in collecting the thermal waters and permitted us to
use their most valuable data and information on the thermal water systems. Most of the surface
waters in Appendix 2 were collected through the courtesy of the City Water Divisions of Ibusuki,
Atami, Ito, Higashi Izu, Shingu, Totsugawa, Kusatsu, Tsumagoi and Kobe and by Mr. M. SUZUKI of,
then, Tokyo Kyoiku University. Our sincere thanks are due to all these people and organizations.
We also thank Misses S. MASHIMA and T. UEMURA and Mrs. M. YAMASAKI of the Institute for
Thermal Spring Research, Okayama University for their skillful assistance in many branches of the
study.
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140 0. MATSUBAYA et ad.
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D/H,180/160 and major element chemistry of thermal waters147
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APPENDIX
5D and 5180 of meteoric2
waters inJapan')
Sample No. Locality Type of water2) 6180,%0 SD, %o
Northeast Honshu
M-1 (73060)3) Odate, Akita-Ken
M-2 (73062) Odate (Shimouchi River)
M-3 (70775) Surface stream, Tamagawa, Akita-Ken
Inner Southwest Honshu
M-4 (73114)
M-5 (73115)
M-6 (73116)
M-7 (73103)
M-8 (72024)
M-9 (72030)
M-10 (73094)
M-11 (73052)
Central Honshu
M-12 (73072)
M-13 (73073)
M-14 (73075)
M-15 (73076)
M-16 (73083)
M-17 (73084)
M-18 (73089)
M-19 (73091)
M-20 (73092)
M-21 (73077)
M-22 (73078)
M-23 (73080)
M-24 (73108)Senami-1 (Suimeiso), Niigata-Ken
Senami-2 (Ryusenkaku)
Senami-3 (Miomote River)
Muikamachi,
Noto-1 (Koiji), Ishikawa-Ken
Noto-2 (Noroshi)
Kanazawa, Ishikawa-Ken
Kurayoshi (Tenjin River), Tottori-Ken
Atami-1, Shizuoka-Ken
Atami-2
Atami-3 (Tanna Tunnel)
Atami-4 (Aizome River)
Ito-l, Shizuoka-Ken
Ito-2
Ito-3
Ito-4
Ito-5
Atagawa-1 (Shirota River), Shizuoka-Ken
Atagawa-2
Kusatsu, Gunma-Ken
Manza, Tsumagoi, Gunma-Ken
Outer southwest Honshyu and Shikoku
M-25 (73069)
M-26 (73096)
M-27 (73098)
M-28 (73099)
M-29 (73105)
M-30 (73106)
M-31 (73071)
M-324) (73110)
M-334) (73214)
M-34 (73113)
M-35 (73228)
M-36 (73229)Shingu (Tenno River), Wakayama-Ken
Shirahama-1, Wakayama-Ken
Shirahama-2 (Mabudani River)
Shirahama-3 (Tonda River)
Nachi-1 (Koza River), Wakayama-Ken
Nachi-2 (Nachi Fall)
Totsukawa, Nara-Ken
Arima-1 (Arima River), Hyogo-Ken
Arima-2 (Arima River)
Arima-3 (Rokko River)
Ishibotoke-1, Kawachi-Nagano, Osaka
Ishibotoke-2 (Amami River)G
R
R
G
G
R
R
G
G
G
R
G
G
G
R
G
G
G
G
G
R
G
G
R
R
G
R
R
R
R
R
R
R
R
G
R 8.9
9.2
-11.0
8.0
8.3
9.0
-10.2
7.9
7.8
8.7
9.1
7.9
7.7
8.0
7.9
7.8
7.4
7.4
6.9
7.7
6.9
6.8
-11.9
-13.0
6.1
5.9
6.0
6.3
5.8
6.4
7.5
8.3
7.7
8.3
7.4
7.9-54.0
-56.3
-65.0
-46.4
-46.3
-48.4
-56.9
-44.2
-42.8
-47.6
-54.0
-47.8
-47.7
-49.5
-49.0
-45.5
-42.9
-45.0
-41.1
-44.9
-40.3
-41.1
-77.4
-84.5
-32.9
-37.1
-35.0
-37.4
-33.6
-35.2
-46.0
-48.3
-47.5
-50.6
-48.2
-52.0
(Continued)
D/H,180/160 and major element chemistry of thermal waters 151
Sample No. Locality Type of water2) 5180,%o SD, O/oo
Central Kyushu
M-37 (73063)
M-38 (73064)
M-39 (73065)
M-40 (73066)
M-4 1. (73067)
M-42 (73135)
M-43 (73253)
M-44 (73263)Beppu, Oita-Ken
lbusuki-1, Kagoshima-Ken
Ibusuki-2
Ibusuki-3
Ibusuki-4
Ibusuki-5 (Ikeda Lake)
Ibusuki-6 (Unagi-Ike)
Ibusuki-7 (Kagami-Ike)R
G
G
G
G
L
L
L8.5
7.2
6.8
3.7
6.9
2.9
3.9
2.6-52.8
-43.4
-41.7
-24.2
-42.0
-19.8
-26.6
-19.4
1)
2)
3)
4)Samples were collected during the period from January to May,
G : ground water; R : river; L : lake.
Figures in parentheses are sample numbers at Misasa.
M-32 collected in January, 1973 and M-33 in May, 1973.1973.
|
Matsubaya (1973) H and O isotopic ratios and major elemtn chemistry of japan thermal systems.txt
|
The Zsland Arc (1994) 3, 182-198
Thematic Article
Tectonic evolution of lower crustal rocks in an exposed magmatic arc
section in the Hidaka metamorphic belt, Hokkaido, northern Japan
TSWOSHI TOYOSHIMA,~ MASAYUKI KOMATSU~ AND TOSU SHIMURA3
'Department of Geology, Faculty of Science, Niigata University, 8050 Zkarashi-2-nocho, Niigata 950-21,
2Department of Earth Sciences, Faculty of Science, Ehime University, 2-5 Bunkyocho, Matsuyama 790 and
3Graduate School of Science and Technology, Niigata University, 8050 Ikarashi-2-nocho, Niigata 950-21, Japan
Abstract The Hidaka metamorphic belt consists of an island-arc assembly of lower to upper crustal
rocks formed during early to middle Paleogene time and exhumed during middle Paleogene to
Miocene time. The tectonic evolution of the belt is divided into four stages, Dors, D,, D,rs, and D,,
based on their characteristic deformation, metamorphism, and igneous activity. The pre-
metamorphic and igneous stage (Do) involves tectonic thickening of an uppermost Cretaceous and
earliest Tertiary accretionary complex, including oceanic materials in the lower part of the complex.
D, is the stage of prograde metamorphism with increasing temperatures at a constant pressure
during an early phase, and with a slight decrease of pressure at the peak metamorphic phase,
accompanying flattening of metamorphic rocks and intrusions of mafic to intermediate igneous
rocks. At the peak, incipient partial melting of pelitic and psammitic gneisses took place in the
amphibolite-granulite facies transition zone, the melt and residuals cutting the foliations formed
by flattening. In the deep crust, large amounts of S-type tonalite magma formed by crustal ana-
texis, intruded into the granulite facies gneiss zone and also into the upper levels of the metamor-
phic sequence during the subsequent stage. During D, stage, mafic and intermediate magmas
supplied and transported heat to form the arc-type crust and at the same time, the magmatic
underplating caused extensional doming of the crust, giving rise to flattening and vertical uplifting
of the crustal rocks. D, stage is characterized by subhorizontal top-to-the-south displacement and
thrusting of lower to upper crustal rocks, forming a basal detachment surface (dkollement) and
duplex structures associated with intrusions of S-type tonalite. Deformation structures and tex-
tures of high-temperature mylonites formed along the dkcollement, as well as the duplex struc-
tures, show that the D, stage movement occurred under a N-S trending compressional tectonic
regime. The depth of intra-crustal dkcollement in the Hidaka belt was defined by the effect of
multiplication of two factors, the fraction of partial melt which increases downward, and the fluid
flux which decreases downward. The crustal dhcollement, however, might have extended to the
crust-mantle boundary and/or to the lithosphere and asthenosphere boundary. The subhorizontal
movement was transitional to a dextral-reverse-slip (dextral transpression) movement accompa-
nied by low-temperature mylonitization with retrograde metamorphism, the stage defined as D,.
The crustal rocks from the basal dkcollement to the upper were tilted eastward on the N-S axis and
exhumed during the D, stage. During D, and D, stages, the intrusion of crustal acidic magmas
enhanced the crustal deformation and exhumation in the compressional and subsequent transpres-
sional tectonic regime.
Key words: deformation history, duplex, Hidaka metamorphic belt, lower crustal rocks, magmatic
activity, magmatic underplating, metamorphic history.
Tectonic evolution of Hidaka lower crust 183
subhorizontal crustal displacement (D2) by Ko-
matsu et al. (1989). The granitic rocks are almost
S-type tonalites formed by anatexis in the deep
crust (Tagiri et al. 1989; Osanai et al, 1991;
Shimura et al. 1992), and intruded into the decolle-
ment and shear zones (Toyoshima 1991; Shimura
1992). Komatsu et al. (1994) showed that granu-
lite facies gneisses underwent decompression under
increasing temperature conditions during the peak
metamorphic and deformation stage (D,) and sug-
gested that the unloading was caused by exten-
sional doming due to intrusion and underplating of
mafic magma. We consider that the mafic rnagma-
tism and the resultant acidic magmatic activity
make a vital contribution to the tectonic evolution
of magmatic arc crust.
This paper is aimed to contribute to the discus-
sion on the role of magmatic intrusions in the
tectonic evolution of the arc crust exposed in the
Hidaka belt. The descriptions are focused on the
foliation-forming deformations with prograde and
retrograde metamorphism and to the fracture-
forming deformations, in which the fractures are
invaded by magmatic injections. All mineral abbre-
viations used are after Kretz (1983). INTRODUCTION
Exposed cross-sections of continental and arc
crusts have attracted the attention of earth scien-
tists investigating the evolutionary processes of the
crust (e.g. Coward et al. 1982, 1986; Percival &
Card 1983; Gibson 1990; Oliver 1990; Percival
1990; Rutter et al. 1993; Salisbury & Fountain,
1990; Treloar et al. 1990). The Hidaka metamor-
phic belt is one example of an exposed cross-
section of magmatic arc. The Hidaka crust was
formed during early to middle Paleogene time and
exhumed during middle Paleogene to Miocene time
(Komatsu et al. 1983, 1989). The general evolu-
tionary history of the belt has been described in
terms of four stages; tectonic thickening of accre-
tionary sedimentary rocks with oceanic materials
and trench-fill materials (D,rs), followed by intrusion
of mafic to intermediate igneous rocks that gave
rise to progressive metamorphism accompanied by
crustal anatexis (Dl), a top-to-the south displace-
ment and thrusting accompanying intrusion of
S-type tonalite (D,) and subsequent dextral-
reverse-slip movement accompanied by strong
mylonitization and retrograde metamorphism (D,;
Komatsu et al. 1989, 1994). The details of defor-
mational aspects of the evolutionary history have
been analyzed by Toyoshima (1991) who divided
the structural history into six phases of foliation-
forming and four phases of fracture-forming defor-
mations during the D, to D, stages. The details of
the metamorphic aspect, especially those of the
prograde P-T path and anatectic reactions in the
granulite zone during D, stage and retrograde P-T
path through D, to D, stages have been analyzed
by Osanai et al. (1991, 1992) and Komatsu et al.
( 1994).
The tectonic and evolutionary history of the
Hidaka metamorphic belt, however, is still contro-
versial, especially with respect to the relationships
between magmatism, metamorphism and deforma-
tion. Jolivet and Miyashita (1985) proposed, mainly
based on the analysis of the parallel western meta-
morphosed ophiolite belt on the western flank of
the Hidaka metamorphic belt, that a right-lateral
strike-slip shear movement was contemporaneous
with the granitic magmatism in the Hidaka meta-
morphic belt, and regarded the resultant crustal-
scale mega-shear movement as a unique movement
ascribed to the tectonic evolution of the Hidaka
metamorphic belt (Jolivet & Huchon 1989). Their
crustal-scale mega-shear corresponds to the de-
tachment surface (decollement) and its extension
to the shallower level of the crust at the time of OUTLINE OF GEOLOGY AND METAMORPHISM
The Hidaka metamorphic belt is bounded on the
west by the Poroshiri ophiolite belt (Western Zone)
in the north and by the late Cretaceous accretion-
ary complex in the south. The boundary is a large
fault (Hidaka Main Thrust) associated with a N-S
trending well-developed mylonite zone (western-
margin mylonite zone) that occupies the western-
most part of the Hidaka metamorphic belt through-
out almost the entire part of the belt. The crustal
metamorphic rocks are exposed with a N-S trend-
ing zonal arrangement (Fig. l), the grade decreas-
ing from granulite facies in the basal part to very
low-grade metasediments on the top of the crustal
metamorphic sequence. The most prominent folia-
tions of the metamorphic rocks are generally par-
allel to lithological boundaries which are generally
subparallel or oblique to the western-margin mylo-
nite zone, and steeply inclined to the east, except in
the southern area. The most prominent mineral
lineations on the foliations are parallel or sub-
parallel to the strike of the foliation in most parts
of the lower sequence, though in some parts,
especially in the southern part of the belt, their
trend changes from a direction parallel to the
strike of the foliation to parallel to the dip of the
14401738, 1994, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1994.tb00106.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
184 T. Toyoshima et al.
Melange-type sediment ra Very low-grade metasedi I I I I I
I I I Uppersequence
1 ' I B Lower sequence
m II/-I Upper gabbro-diorite
D3 tonalite-quartz diorite
Lithologic boundary
_--
---
0 iment
Fig. 1
Hidaka Simplified geological map of the
metamorphic belt (HMB).
foliation. The upper (eastern most) metasediments
are lithologically continuous with the sedimentary
rocks of Nakanogawa Group (earliest Tertiary,
Nanayama 1992) overlying the metamorphic se-
quence (Komatsu et al. 1983). The maximum total
thickness of metamorphic rocks plus overlying sed-
imentary rocks is estimated to be about 23 km
(Komatsu et al. 1983, 1989).
In the metamorphic sequence, metamorphosed
mafic rocks are dominant in the lower (western)
part, and pelitic and psammitic metamorphic rocks
are dominant in the upper (eastern) part. The former is called the lower sequence, the latter is called the
upper sequence (Komatsu et al. 1983). The upper
sequence consists, in descending order, of prehnite-
pumpellyite grade metasediments, chlorite (Ch1)-
muscovite (Ms)-biotite (Bt) schists (greenschist
facies), Ms-Bt schists and gneisses, and Bt gneisses.
The lower sequence consists of amphibolites and Bt-
garnet (Grt) gneisses (amphibolite facies), Bt-Grt-
cordierite (Crd) gneisses, Grt-Crd-orthopyroxene
(Opx) gneisses, and Opx-amphibolites (granulite
facies; Komatsu et al. 1983; Osanai 1985; Osanai
et al. 1986a). The highest P-T conditions are esti-
14401738, 1994, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1994.tb00106.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Tectonic evolution of Hidaka lower crust 185
were formed through D, to D, stages and post-D,
phases of tectonic history (Toyoshima 1991). The
deformation style in the tectonic history changes cy-
clically from ductile deformation (folding, shearing
and formation of foliation) to brittle deformation
(fracturing and faulting). The deformations in the
early phases are of the penetrative type, and in the
later phases are of non-penetrative type. The later
deformations are mainly concentrated in localized
shear zones along the western boundary fault and
thrusts. The latest deformation phases (post-D3) are
non-penetrative, mostly concentrated in sinistral
shear zones and truncated faults whose formation
can be related to the oblique subduction of the Pa-
cific plate (Toyoshima 1990), and therefore for the
purpose of this paper they can be ignored.
During the exhumation movement (D, and D,) of
the crustal metamorphic rocks, many mylonite zones
were formed at different crustal levels, most of them
converging into the basal mylonite zone (western-
margin mylonite zone; Komatsu & Toyoshima
1992). The basal mylonite zone, which is occupied by
mylonitized tonalites, is graded structurally upward
into undeformed or very weakly deformed tonalites,
followed by a high-grade gneiss unit (granulite unit)
in most parts of the metamorphic belt. Deformations
during early phases of the exhumation produced
N-S to NW-SE trending mylonitic foliations and
mineral lineations of horizontal or gentle northward
plungings. The early structures are rotated, folded,
and modified in other fashions through later-phases
deformations.
In this paper, we redefine the deformation
phases from those proposed by Toyoshima (1991)
into three foliation-forming phases and two inter-
phases forming fractures filled with igneous veins
which are contemporaneous with intrusives occur-
ring on geological map scales (Fig. 1; Table 1). We
will describe in the following section the character-
istic rock structures and mineral textures produced
in individual stages (D,-D,), in order to examine
the relationship between deformation, metamor-
phism, and igneous activity during the evolutionary
history of the Hidaka metamorphic belt. mated to be ca 7 kb, 750°C during the prograde
stage, attending thermal peak conditions up to
800°C at ca 6 kb (Komatsu et al. 1982,1989,1994;
Osanai 1985; Osanai et al. 1986a,b). The regional
metamorphic gradient at peak conditions is attained
at ca 40 ' C/km. A complete metamorphic sequence
is found in the Shizunai River region in the central
part of the metamorphic belt (Osanai 1985). In other
parts the sequence is disturbed by large intrusive
bodies and by several sets of top-to-the-south
thrusts and dextral strike-slip faults. In the central-
southern part, metamorphic rocks are duplicated by
these faults, and a low-angle pile of nappes of meta-
morphic rocks has been formed (Arita et al. 1986;
Owada 1989).
The metamorphic rocks are intruded by a large
amount of mafic to acidic plutonic rock (Fig. 1)
whose intrusive phases range from the time of D,
through D,, to D,-deformations. Early gabbroic to
dioritic intrusives, which are partly metamorphosed
and mylonitized under granulite facies conditions,
form sheet-like or lenticular bodies, and are em-
placed subparallel or oblique to the lithologic
boundaries with the metamorphic country rocks.
The acidic intrusives, most of which are peralumi-
nous (S-type) and show sheet-like or lenticular
shapes, have been classified into four depth types;
upper granite-granodiorite, middle (3rd-Ms to-
nalite, lower Crd-Grt tonalite, and basal Grt-Opx
tonalite (Komatsu et al. 1986, 1989; Shimura et
al. 1992). The granitic rocks include a large
amount of country metamorphic rocks that they
are emplaced in. The difference of the granitic
rocks was caused by a crustal density filter mech-
anism (e.g. Glazner & Ussler 1988) that ponds
magmas at different structural levels according to
their composition, density, temperature and viscos-
ity, and by the difference of crystallization condi-
tions at different structural levels. Therefore the
granitic magmas were emplaced in a flat-lying
structural state into the Hidaka crust (Komatsu
1986; Shimura 1992). The basal tonalitic intrusive
zone is largely mylonitized and forms the western-
margin mylonite zone. Spinel and plagioclase lher-
zolite bodies are also found within the basal part of
the belt (Komatsu et al. 1983; Niida 1984).
DEFORMATION HISTORY
Seventeen phases of deformation are recognized in
the evolutionary history of the Hidaka metamorphic
belt on the basis of structural and petrological analy-
sis of rock structures and mineral textures which PROGRADE METAMORPHIC PEAK ACCOMPANYING
FLATTENING AND MAFIC INTRUSION (DI-STAGE)
S, -flattening f o I I owed by mafi c intrusion
The &-phase of deformation is defined by bedding
foliations and the formation of boudins (S,-boudins)
of quartzo-feldspathic veins and thin layers of
coarse-grained amphibolite intercalated within fine-
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186 T. Toyoshima et al.
Table 1 Tectonic and metamorphic history of the lower sequence of the Hidaka metamorphic belt. Relative temperature is
metamorphic temperature conditions of the granulite unit.
Deformation & structure Igneous activity
? ?
Flattening & shearing
Bedding foliation (Sl) & Gabbrdiorite
non-rotated boudinage
Locally preserved fold
with axial foliation (SI')
Shearing, folding & Tonalite-quartz
mylonitization diorite intrusion
with axial foliation (52)
Duplex thrust of duplex Mylonitic foliation &fold into decollement
8, ramwoof eformatioi-
age Metamorphism
Low-grade
metamorphism
Prograde
metamorphism &
anatexis
Metamorphic peak
up to granulite facie!
Initiation of
metamorphism retrograde DO
D1
D2
D3 Tectonics
Tectonic
thickening of
sedimentary precursors
Magmatic
underplating
& extensional
doming
Initiation of
decollement
Subhorizontal
top-to-the
south displacement
Eastward
tilting
&
dextral-slip
movement Shearing, folding & Retrograde metamorphism
retrograde mylonitization 1 Minor intrusion 1 lower Dart: aranulite to
of olivine greenschist facies
greenschist facies upper part: amphibolite 1 Mylonitic foliation &
fold with axial foliation (S3) Relative
temperature f---
?
I/i itructural ,tate of
,hear plane
?
Flat-lying
Flat-lying
Steeply
inclined
grained amphibolites in the lower metamorphic se-
quence. The S,-boudins have a pancake shape, flat-
tened parallel to the S,-foliations and to the
lithological bandings (Fig. 2a). No rotated boudins
of S,-phase are observed. Thus, it can be said that
the S,-deformation is of the flattening type. In am-
phibolites that escaped post-S,-phase ductile defor-
mation, no mineral lineation is visible in the S,-
foliations, or, they are very weak. The preferred
lattice fabric of brown hornblendes in the amphibo-
lites has an axial symmetry (Fig. 3). The fabric pat-
tern indicates that the crystallographic c-axes and
(100) planes of the brown hornblendes are randomly oriented on the S,-foliation and a great-circle girdle
is defined by the distribution of the c- and b-axes.
At the edges of S,-boudins, internal foliations
are curved and necked, and pyroxene or brown
hornblende grains are also curved along the outer
surfaces of the boudins. Separations of S,-boudins
in granulite facies gneisses are commonly filled
with leucocratic aggregates of coarse-grained or-
thopyroxene or brown hornblende. In the neck
regions of S,-boudins in amphibolite facies gneis-
ses, coarse grains of brown hornblende forming the
boudins are partly replaced by fine grains of brown
hornblende. Most of these grains of pyroxene and
Fig. 2 Photographs of S,-boudins in pyroxene arnphibolite. (a) Coarse-grained amphibolite layers (C) separated into sausage-shaped boudins not rotated, (b) D2
stage fold of S,-foliation with S,-boudins.
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Tectonic evolution of Hidaka lower crust 187
L L
L
X
m] 5.4-7.6%
2.6-4.2% w 7.4-9.0%
m] 4.2-5.8% N=100 X 0 1.1-4.0°/o ~8.0-10.0%%
[IIII 6.0-8.0% 4.0-6.0°/o 10.0-1 2.0%
Y
6.0-8.0%
3.6-5.2% 8.4-1 0.0%
rrnr 5.2-6.8% N=lOO X
5.4-7.6%
[IIII 6.6-9.4%
4.0-7.0% 13.0-1 6.0%
[rm 7.0-1 0.0% N=lOO
Fig. 3 Change in preferred lattice fabric of brown hornblendes from &deformed amphibolite to D,-amphibolite mylonite
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188 T. Toyoshima et al.
hornblende do not or, if any, very weakly show
undulose extinction, kink bands, and deformation
twins. These facts indicate that the orthopyroxene
in the granulite facies gneisses and the brown
hornblende in the amphibolite facies gneisses were
stable during the S,-phase, and crystallized just
before or during the S,-phase, and were annealed
after the S,-phase. This suggests that the meta-
morphic grade during the &-phase approached the
climax conditions of the prograde metamorphism
of the Hidaka lower crustal rocks.
Shear deformation (S, ’-phase)
The deformation of the S, ’ -phase is characterized
by the formation of S,-parallel mylonitic foliations
(S,’) and by folding of the S,-foliations into isocli-
nal shapes with well developed axial foliations
(Fig. 4). The S,’-foliation of psammitic gneisses or
amphibolites is defined by elongated gedrites
and/or hornblendes and those of pelitic gneisses by
biotites. The S,’-foliations are cut by small patches
or veins that are made up of randomly oriented
coarse-grained minerals without hydrous phases
except for cordierite. From mineral assemblages
and textures, the small patches or veins are con-
sidered to be anatectic melts formed at peak meta-
morphic conditions (Komatsu et al. 1994). The
metamorphic minerals defining the S,’-foliations
suggest that the metamorphic temperature during
the S,’-phase was slightly lower than that of the
Ged-PI-(Opx) PI-Qtz-(Opx-Crd) Opx-Crd-(PI) .‘ thermal peak of the prograde metamorphism of the
Hidaka lower crustal rocks (Komatsu et al. 1994).
Mafic intrusion concurrent with the highest grade
metamorphism
The S,- and S,’-foliations are cut by mafic veins.
These fractures in the granulite unit are intruded by
pyroxene gabbroic to dioritic rocks, whereas those in
the amphibolite facies gneiss unit (brown horn-
blende amphibolite unit in the middle level of the
lower metamorphic sequence) and low-grade am-
phibolite facies gneisses (gneisses in upper level of
the lower sequence) are intruded by hornblende gab-
broic to quartz dioritic rocks. The difference in the
mineral assemblages implies that the mafic intrusion
was emplaced in a flat-lying structural state of the
Hidaka crust. Most of the mafic intrusive rocks have
S,-foliations, and have recrystallized under granu-
lite facies conditions in the granulite facies gneiss
unit and under amphibolite facies conditions in the
amphibolite facies gneiss units, respectively. There-
fore the &-phase deformation took place just after
the intrusion of mafic rocks, during the time that
high temperature metamorphic conditions were still
maintained.
The main bodies of gabbroic to dioritic rocks occur
as large lenticular masses in the northern and south-
ern parts of the belt (Fig. 1) and as intrusive sheets
emplaced in the upper level of the lower sequence
throughout almost the entire part of the belt. The
Fig. 4. Sketch of anatectic melt (Pl-
Qtz-[Opx-Crd]) and residual (Opx-Crd-
[PI]) cutting tight fold with S,’ axial foli-
ation in gedrite gneiss (Ged-PI-[Opx]) of
the granulite unit.Mineral abbreviations
following Kretz (1983).
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Tectonic evolution of Hidaka lower crust 189
TOP-TO-THE-SOUTH SUBHORIZONTAL MOVEMENT
ASSOCIATED WITH HIGH-TEMPERATURE MYLONlTlZATlON
AND S-TYPE TONALlTlC INTRUSION (D,-STAGE) latter intrusives were metamorphosed to Bt-Hbl
gneisses and occupy most of the upper level, which is
therefore called the Bt-Hbl gneiss unit.
Large cordierite porphyroblasts are rarely found,
growing across S,- or S,’-foliations in the matrix of
pelitic gneisses of the granulite unit, which were
deformed during the S,- and S,’-phases but only
slightly deformed in ductile fashion after the intru-
sion of mafic rocks (Fig. 5a). The cordierite grains
show rounded shapes, enclosing fine grains of her-
cynite, ilmenite, biotite and fibrolite. The cordierite
porphyroblasts replace locally or completely large
garnet grains that include fibrolite and small biotite
grains. The cordierites and their inclusion minerals
do not have preferred dimensional and lattice orien-
tations, except for those in the pelitic gneisses that
were strongly deformed in ductile fashion during the
later phases. Such a mode of occurrence and the
fabrics of the cordierite porphyroblasts indicate that
they were produced under non-deformational condi-
tions during the peak metamorphic time. Komatsu
et al. (1994) argued that the gabbroic magma was
intruded into lower metamorphic rocks concurrently
with the anatexis of the pelitic and psammitic
gneisses and that the spinel-bearing cordierite was
formed through the anatectic reactions shown be-
low:
Bt + Grt + Sill + vapour = Crd + Spl + Ilm + melt
Bt + Grt + Qz + vapour = Opx + Crd + Ilm + melt
The anatectic reactions took place at the highest
temperature during the prograde metamorphism
that took a path of increasing temperature with a
slight decrease of pressure (Komatsu et al. 1994). High-temperature mylonitization
The deformation of D, stage is characterized by
the formation of high-temperature S,-parallel my-
lonitic foliation (S2), especially in the mafic intru-
sive rocks, and by the folding of the S, or Slr-
foliations in isoclinal or more rarely open shapes
with well developed axial foliations (S,; Fig. 2b).
Structures of the &-phase are the most prominent
structures in the lower sequence. The S,-boudins
are rotated sinistrally against the S,-mylonitic
foliation. The S,-deformation caused widespread
transposition of layering on various scales. Most of
the folds are of an asymmetrical type with a
dextral sense of shearing (Fig. 6a). Some folds
show a sheath shape and some are of a rootless
intrafolial type, implying a high-strain magnitude
during the D,-folding. The S,-structures are fre-
quently cut across by S-type tonalitic dykes or
veins (Fig. 7).
D,-mylonites show various kinds of asymmetri-
cal rock structures and mineral textures such as
sinistrally rotated boudins, S-C structures (Lister &
Snoke 1984), porphyroclasts with asymmetrical
tails of recrystallized matrix minerals, en echelon
deformation lamellae in orthopyroxene porphyro-
clasts and kinked porphyroclasts (Fig. 6b-f). The
C-surfaces correspond with S,-foliations and are
subparallel to the general trend of the lower se-
quence. These structures and textures indicate that
the deformation during the D, stage was caused by
a dextral sense of shear (Toyoshima 1991).
Fig. 5 Cordierite porphyroblasts (Crd) in pelitic gneiss of the granulite unit, enclosing fine grains of hercynite, ilmenite, biotite and fibrolite. Scale bar is 2.0 mm,
plane-polarized light (PPL). (a) Cordierite porphyroblasts cutting across S,-foliation in matrix, showing rounded shapes and replacing locally large garnet grains,
(b) elongated cordierite porphyroblasts elongated parallel to the S,-foliation.
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190 T. Toyoshirna et al.
Fig. 6 Deformation structures and textures of D2 deformed rocks indicating a dextral sense of shear. (a) Asymmetrical fold with &-axial foliation in gabbro
mylonite. (b) S,-boudins rotated sinistrally (RB) by D2 deformation in the brown hornblende amphibolite unit. Scale bar is 30 cm. (c-f) S-C-(C’) mylonites derived
from D, gabbro, pyroxene gneiss of the granulite unit, from amphibolite of the hornblende amphibolite unit, D, quartz diorite, respectively. Opxpc: orthopyroxene
porphyroclasts, Hblpc: hornblende porphyroclast, Bt: biotite, Hbl: hornblende, Plpc: plagioclase porphyroclast. Scale bars of (c-1) are 2.0 mm, PPL.
The spinel-bearing cordierite grains produced
during the peak metamorphic phase are elongated
parallel or subparallel to S,-foliation in pelitic
gneisses of the granulite unit which were deformed
during D, stage (Fig. 5b). The elongated cordierite
grains define a strong mineral lineation on the
S,-foliation. They are partly or completely replaced
by aggregates of small grains and subgrains of
neocordierite during the D,-deformation. They
show wavy extinction, and the inclusion minerals,
hercynite, ilmenite, biotite and fibrolite, are ar- ranged parallel to the long axes of the elongated
cordierite grains, forming Si-foiiation which is
parallel to the external S,-foliation. The preferred
dimensional orientzition of inclusion minerals is
intensified with increase of elongation of the host
cordierite grains. The preferred orientation is re-
lated to the intracrystalline deformation of the host
crystals.
Amphibolite mylonites produced by the S,-defor-
mation are characterized by the preferred dimen-
sional orientation of elongated brown hornblende
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Tectonic evolution of Hidaka lower crust 191
It is noteworthy that the recrystallized mineral
assemblages in the D,-mylonite zones are almost the
same in each unit as those formed during the D,
stage (Table 2). The D,-mylonite zones are subpar-
allel to the metamorphic layers, hence the mylonite
zones were formed subhorizontally at different
crustal levels. It implies that the D,-deformation
took place in a flat-lying structural state of the crust
and that the apparent dextral-sense of shear move-
ment indicated by microstructures and textures is
originally top-to-the-south subhorizontal displace-
ment of detached crustal rocks.
Deformation
stages
D1
D2
D3 grains (Fig. 6e), which define the L,-lineation and
S,-foliation. The lattice fabric pattern of the brown
hornblendes indicates that the crystallographic
c-axes and (100) planes of the brown hornblende
grains are preferentially oriented slightly oblique to
the L,-lineation and to the S,-foliation (Fig. 3). The
preferred lattice orientation results from the modifi-
cation of fabric pattern with axial symmetry of
brown hornblende grains formed during the D,
stage (Fig. 3). Recrystallized fine-grains of brown
hornblende occur at the tips of coarse brown horn-
blende grains and on their grain boundaries. It is
inferred from the texture and fabric that the coarse
brown hornblende grains were partly replaced by
recrystallized fine grains of hornblende during the
D,-dextral shearing to form elongate shapes and
preferred dimensional orientation defined as L,.
In the D,-mylonites, the mineral assemblage and
mineral chemistry of the matrix minerals are nearly
the same as those of the porphyroclastic minerals
(Toyoshima 1989,1991). Orthopyroxene and garnet
porphyroclasts in D,-mylonites of granulite unit
show only weak chemical zoning. These indicate
that the metamorphic conditions of the Hidaka lower
crustal rocks during D, stage were nearly the same
as those of the highest conditions of temperature in
the Hidaka prograde metamorphism.
Products Granulite unit 8, basal tonalite Arnphibolite unit Biotite-Hornblende gneiss unit
Pelitic or psammitic rocks Basic rocks Basic rocks Pelitic rocks Basic rocks
Grt-Bt-Qtz-Plz*Kfs Hbl-Opx-PI*Qtz Hbl-PI+Spn
S1 &S1’ Grt-Ged-PCOtz Hbl-Pl*Qtz Hbl-PIeQtz ? ?
................................................. Opx-PI+Qtz ..................................................................................................................................................................................................................................................
Opx-Crd
Melt Opx-PI-Qtz-(Crd) ? ? ?
Grt-Crd-Opx-PI-Qtz+Kfs Hbl-PItQtz Hbl-PI+Spn Bt-PI-Qtz greenish brown Hbl-Bt-PI-Qtz+Spr
OPX-Grt-PI-QtZ+KfS HbCOpx-PI+Qtz HbCPCQtz+Cum greenish brown HbCPCQtz+Spn
S2 Opx-Crd-PCQtztKfs Opx-PI+Qtz
Hbl-Cum-PI-Otl
Opx-Cpx-PI*Qtz
I OPX-Bt-QtZ-PI*KfS Opx-PI*Qtz HbCPI+Qtz*Spn Bt-PI-Qtz green Hbl-Bt-PI-Qtz
i Grt-Bt-PI-Qtz+KfS Cum-PI
f Bt-PI-Qtz+Kfs Cum-Bt-PI green Hbl-Pl+Otz Bt-PI-Qtz Act-Bt-PI-Otz
f 2 Bt-Gt-PI-Qtz+Kfs HbCPCQtz*Spn actinolitic Hbl-PI+Qtz Bt-Ms-PI-Qtz
I MS-Chl-PI-Otz Act-PI-Otl Act-PltOtz Ms-Chl-PI-& ChCEp-PCQtz-
I Ms-Chl-Ep-PI-Qtz Act-Chl-Ep-PI-Qtz-Spn
I 3 Ms-Chl-PI-Qtz&al ChCEp-PCQtz+Cal
ChCEp-PCOtz-Spn
ChCMs-Ep-PCQtz+Spn I 1 Grt-Crd-Bt-PI-OtztKfs Hbl-PltQtz
Ath-Pl+Qtz ................................................................................................................................................................................................................................................................
s3* i Hbl-PI-Qtz*Bt ...................................................................................................................................................................................................................................................... Top-to-the-south subhorizontal displacement
promoted by S-type tonalitic intrusion
The fractures invaded by S-type tonalitic to quartz
dioritic intrusives sharply cut across various kinds of
pre-existing structures such as Sl’- and S,-folia-
tions (Figs 7a,c,d). The S-type granitic intrusive
rocks contain large amounts of various-sized meta-
morphic inclusions that preserve pre-existing struc-
tures (Figs 7b,8a). Most of the granitic intrusive
rocks have planar and linear structures, showing
that they, together with surrounding metamorphic
rocks, were deformed after the intrusion (Figs 7b,
8b).
Table 2 Mineral assemblages in pelitic-psammitic rocks and basic rocks of the lower sequence. Mineral abbreviations following
Kretz (1983). Grt: garnet, Crd: cordierite, P1: plagioclase, Qtz: quartz, Kfs: potash feldspar, Opx: orthopyroxene, Bt: biolite, Ms:
muscovite, Chl: chlorite, Ep: epidote, Cal: calcite, Cpx: clinopyroxene, Hbl: hornblende, Cum: cummingtonite, Ged: gedrite, Ath:
anthophyllite, Act: actinolite, Sph: sphene.
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192 T. Toyoshirna et al.
10 cm
(c) -"., Ton
-\
Fig. 7 Sketches of D, stage fold with &-axial foliations, D, stage folds with
S,-axial foliation and D, tonalitic vein (Ton). (a) 0, stage fold cut across by
tonalitic vein. (b) Tonalitic rocks cutting across 0, stage fold with axial
foliation (S,) and parallel arrangement (L,) of minerals (hypersthene and
hornblende) in the D, tonalitic intrusive rocks. (c) D, tonalitic veins cutting
&-axial foliation of D, stage fold but cut across by &-axial foliation (S,) of 0,
stage fold. (d) As above. D2 stage fold nearly coaxial with the D, stage fold.
The D,-folding induced the effect to decrease interlimb angle of the D2 stage
fold. Lens cap of camera in figure (c) is 5.4 cm in diameter. The main bodies of the acidic intrusive rocks are
considered to have been intruded into subhorizontal
shear zones, and develop zonally as two giant
sheets; one along the lowest border (western end)
of the Hidaka metamorphic belt and the other
along the boundary between the lower and upper
sequences (Fig. 1). The former is the basal tonalite
and the latter is the middle tonalite after Komatsu
et aL. (1986). The two intrusive zones correspond
to specific structural breaks in the northern part of
the belt; a basal dkcollement and ramp to roof
thrust of duplexed metamorphic rocks. The dkcol-
lement and duplex-forming tectonics have been
considered to be caused by the subhorizontal dis-
placement of the Hidaka island arc crust (Shimura
1992). The acidic intrusives are also commonly
found as dykes and veins in the granulite unit
forming agmatitic structures.
The tonalitic rocks commonly occur also as
network-like injection veins within basic and pelitic
gneisses, extending from thick tonalite dykes. In
some outcrops, the injection veins filled a set of
Riedel shear fractures, corresponding to composite
planar fabrics, S (Berth6 et aL. 1979; Lister &
Snoke 1984) and C' surfaces (Ponce de lon &
Choukroune 1980; Fig. 9), showing that the frac-
tures were formed as layer-parallel dextral shear-
ing, namely top-to-the-south displacement.
RETROGRADE LOW-TEMPERATURE MYLONlTlZATlON
RESULTED FROM DEXTRAL-REVERSE-SLIP MOVEMENT
(D,-STAGE)
The deformation of the D, stage is characterized
by the formation of mylonite zones including the
western-margin mylonite zone, accompanying ret-
rograde metamorphism. As the deformation pro-
Fig. 8 Shapes of metamorphic inclusions (Inc) in D2 tonalitic rocks (Ton). (a) Weakly deformed D, tonalitic rock including weakly elongated metamorphic
inclusions. (b) Strongly deformed tonalitic rock during D, deformation including rolated and stretched metamorphic inclusions.
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Tectonic evolution of Hidaka lower crust 193
tization also developed folding of pre-S,rs-foliations
in tight to isoclinal shapes with well developed axial foli-
ations (S3). Some folds are of asymmetric type with
Z-shaped profiles, and some show mushroom, sheath
and rootless forms, owing to a high-strain magni-
tude during the folding. In many places the folding
progressively induced the effect to decrease inter-
limb angles of the folds (Fig. 7d).
The mylonites show asymmetrical rock structures
and mineral textures such as asymmetrical folds,
sinistrally rotated boudins, S-C-C’ (Fig. 10a) or
S-C (fig. 11) structures, rotated porphyroclasts
with asymmetrical tails of recrystallized minerals
(Fig. lob), and kinked porphyroclasts. These struc-
tures and textures indicate that the S,-mylonite-
forming deformation took place due to shearing in a
dextral sense (Toyoshima 1991).
In those rocks that suffered the mylonitization of
earlier phases, retrograde metamorphic minerals
produced along the S,-foliation are different among
the three units of the lower sequence (Table 2). In
the granulite unit, the recrystallized mineral assem-
blages indicate that metamorphic temperature de-
creased down to that of low granulite facies. On the
other hand, only a slight change of mineral assem-
blage is detected in the two amphibolite facies gneiss
units (brown hornblende amphibolite and Bt-Hbl
gneiss units), showing that the temperature was
almost constant throughout the deformation stages
from D, to the early phase of D, in these units. The
decrease of P-T conditions of the basal unit implies
that the unit exhumed to nearly the same level as the
middle and upper units, and that the Hidaka crustal
rocks, which were in a flat-lying structural state
before the D,-stage, were tilted eastward and
rotated around horizontal N-S axes during the D,-
stage. Therefore, E-W compressional sense of
movement was added to the dextral strike-slip NlO’W I
45ow
5cm
1 Ocm -
7
Fig. 9. Sketches showing tonalite veins (spotted) injected into banded
biotite gneiss (striped). (a) Injection along asymmetric fold axis, (b) injection
along a set of Riedel shear fractures, corresponding to S and C’ surfaces,
formed by the dextral shear movement.
gressed, the mylonitization was locally concentrated
to narrow zones and to the western mylonite zone in
the later phase. The mylonitization developed folia-
tion (S,) with a mineral lineation (L3) which is de-
fined by retrograde metamorphic minerals recrystal-
lized or elongated (Table 2). The S,-structures are
prominent in the lower sequence. The acidic intru-
sive rocks first suffered the mylonitization, together
with their country rocks. Metamorphic inclusions in
the acidic intrusive rocks are sinistrally rotated and
stretched by D,-deformation (Fig. 8b). The myloni-
Fig. 10 Microstructures and textures of D3 deformed rocks. Scale bars are 2.0 mm. (a) S-C-C’ mylonite derived from D2 tonalite. PPL. (b) Garnet grains with
dimensional oriented inclusions (Bt) and asymmetrical tails (Btt) of biotite. Crossed-polarized light.
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194 T. Toyoshima et al.
Fig. 11. Biotite grains (Bt) occurring as wings of large garnet grains (Grt) in
D3 mylonite. Composite planar fabrics in the form of C- and S-surfaces are
found. Scale bar in 2.0 mm, PPL.
movement, namely, dextral transpression acted
since D,-stage.
In the later phase of the D, stage, the deformation
is localized in narrow zones forming strong mylonite
zones, and to the western mylonite zone. Recrystal-
lized mineral assemblages during the later phase are
nearly the same among the three units of the lower
sequence (Table 2). In the granulite unit and basal
tonalite zone, neo-crystallized hornblende grains re-
place hornblende formed in the earlier phase, which
are lower Na and Ti contents but higher Si and A1
contents than the latter, and indicate that the meta-
morphic conditions in the granulite unit and basal
tonalite zone are just below or close to the lower
stability limit of brown hornblende, whereas in the
brown hornblende amphibolite and Bt-Hbl gneiss
units, recrystallized hornblende is brownish-green
to green in color, the condition being below the lower
stability limit of brown hornblende. The change in
metamorphic condition is not so large in the upper
two units as in the lower unit, which implies that the
upper units cooled with extreme slowness through-
out the D, stage. On the other hand, the metamor-
phic conditions of the lower zones of granulite unit
and basal tonalite declined to the same metamorphic
grade (low-grade amphibolite facies) as those of the
upper two units during the later phase.
In the western-margin mylonite zone, olivine
gabbros or hornblende diorite dykes (with varied
thickness from 5 cm to 20 m) are intruded, gently obliquely cutting the mylonite foliation. These
dykes include metamorphic rocks that underwent
the D,-deformation of later phase, and the dykes
themselves show weak mylonitic foliations with
retrograde minerals such as chlorite and muscovite.
In the latest phase, the deformation is further lo-
calized to the western border of the western-
margin mylonite zone and also to very narrow shear
zones with less than 10 em wide in the lower meta-
morphic sequence. The western-margin mylonites
show a subvertical mylonitic foliation with a subhor-
izontal strong mineral lineation. The resultant tex-
tures of the western-margin mylonites change grad-
ually toward the west from coarse porphyroclastic to
phyllitic and the amounts of porphyroclasts de-
crease, but the proportion of recrystallized fine-
grained minerals increase. Microstructural evidence
indicates that the strain magnitude of the myloniti-
zation increased progressively toward the west dur-
ing the latest phase. The gradual change of mylonite
textures is in harmony with the change of retro-
grade metamorphic grade toward the west from
lower amphibolite facies to greenschist facies
(Table 2).
Narrow mylonite zones produced within the
lower sequence dip gently to the east or northeast,
but some of them are nearly horizontal, cutting
obliquely at high angles the previous mylonite
foliation. Mineral lineation plunges very gently
toward the north. These narrow mylonite zones
have greenschist facies mineral assemblages, the
same as those of the mylonites of the western part
of the western-margin mylonite zone. At the later
phase of D, stage, the whole of the lower sequence
was eventually emplaced under the same metamor-
phic conditions. It can be concluded that the D,
stage movement gave rise to the transformation of
structural state from the flat-lying structural state
to the steeply eastward inclined structural state.
ROLE OF MAGMATIC INTRUSION AND RHEOLOGICAL
STRATIFICATION: DISCUSSION
In the previous sections, we have described the de-
formational and metamorphic aspects of the evolu-
tional history in the lower sequence of the Hidaka
metamorphic belt, focusing on foliation-deformation
and igneous activities during D, through D,. The
evolutionary history and tectonics are summarized
in Table 1. The igneous intrusives took part in these
events, particularly gabbroic and dioritic intrusives
in D,, and S-type tonalite intrusives in D,.
The P-T-D path of the lower metamorphic se-
quence shows that during the prograde metamor-
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phic stage, the Hidaka crustal rocks were affected
by flattening and shearing and attained the highest
temperature conditions, accompanied by a slight
decrease of pressure, due to unloading (Fig. 12).
The fact that metamorphic temperatures increased
during the unloading implies that the unloading
was caused by magmatic underplating (cf. Rutter
et al. 1993) beneath the crust.
At peak metamorphic conditions, incipient partial
melting of pelitic metamorphic rocks took place in
the lowest part of the lower metamorphic sequence,
which corresponds with the depth of upper part of
the lower crust (Osanai et al. 1991; Komatsu et al.
1994). The products of partial melting cut Sl'-
foliations in the pelitic gneisses, and the mineral
products have no preferred lattice orientation (Ko-
matsu et al. 1994). The D,-gabbros also cut Sl'-
foliations and the gabbros themselves do not have
any deformed structure, except for the rocks that
suffered later mylonitization. Therefore, in the upper
part of the lower crust, the formation of fractures,
invasion of gabbroic magmas and partial melting of
pelitic gneisses occurred during the same time inter-
val as the D, stage in a tectonic setting free from
compressional and shear stress. It should be noted
that in the upper part of the lower crust, the mafic
magmatic intrusion was emplaced subsequent to the
flattening of the crustal rocks that accompanied pro-
grade metamorphism and crustal anatexis, caused
by underplating of the mafic magma. It implies that
the thermal pe8.k at any level of the crust was at-
tained at the time when the magma reached that
level.
In the deeper part of the crust, an S-type tonalite
magma was produced through extensive crustal
anatexis during the D,-magmatic intrusion (Owada
1989; Tagiri et al. 1989; Osanai et al. 1991;
Table l), and emplaced along the basal dbcollement,
kb
Gabbro-diorite
Tonalite-quartz
diorite intrusion
6bO 8bO "C
Temperature 460
Fig. 12 The P-T-D history of the lowest part of the lower sequence, Hidaka
metamorphic belt (partly modified from Komatsu et a/. [1989] and Toyo-
shirna [1991]). Triple point of aluminosilicates after Salje (1986). Tectonic evolution of Hidaka lower crust 195
ramps and roof thrusts of duplexed metamorphic
rocks (Shimura 1992), especially along the dkcolle-
ment and roof thrusts where giant planar intrusives
are formed (Toyoshima 1991). The dbcollement was
formed along the level where the anatexis of crustal
rocks is initiated and intensified downward. This
clearly illustrates where a crustal scale dbcollement
forms in a magmatic arc in a compressional tectonic
regime.
The incipient partial melting of pelitic and psam-
mitic gneisses of the belt occurred above the dkcol-
lement at around 800"C, 6 kbar (Komatsu et al.
1994) under water-saturated or fluid-present condi-
tions (Osanai et al. 1992), and the extensive produc-
tion of S-type tonalite magmas is supposed to have
occurred at above 900°C, > 8 kbar under H,O-
undersaturated or fluid-absent conditions (Shimura
et al. 1992). It is expected that the melt fraction was
significantly different above and below the dbcolle-
ment. According to the experimental work on gra-
nitic rocks under partial melting conditions (Dell'
Angelo & Tullis 1988), partially melted rocks remain
strong for smaller amounts of melt, but they show
generally a significant decrease in strength for melt
fractions of more than 15%. A high fluid pressure,
under which water-saturated partial melting was
promoted, is also expected to weaken the rocks, as it
allows fractures to form (Dell'Angelo & Tullis
1988). Therefore, a mechanically weak level of the
crust is defined by the fraction of partial melt which
increases downward and by fluid flux which de-
creases downward. It can be said that the presently
exposed part of the lower crust was relatively
stronger and more brittle than the non-exposed
lower part in the Hidaka crust, depending on the
melt fraction. It has been argued that the weak
zones on one side of strength discontinuities will be
preferred sites for the development of major, sub-
horizontal detachment faults (dbcollement; e.g.
White & Bretan 1985; Ord & Hobbs 1989). Crustal
anatexis in the Hidaka crust may have determined
the depth of the dkcollement and the maximum
depth exposed in the Hidaka metamorphic belt.
More important rheological discontinuities in the
magmatic arc are the boundaries between the crust
and mantle (Moho) and between lithosphere and as-
thenosphere. Mantle detachments, exposing Moho,
are inferred in the Ivrea (Berckhemer 1969) and
Kohistan (Coward et al. 1986) sections. The depth
of detachment, which depends on crust and mantle
rheology at the time of deformation, determines the
maximum depth exposed (Percival et al. 1992). De-
tachment at the Moho and/or between the litho-
sphere and asthenosphere may have been active in
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196 T. Toyoshirna et al.
the Hidaka metamorphic belt before crustal melting,
because upper-mantle peridotites, most of which are
mylonitized, occur in abundance along the dkcol-
lement of the belt (Komatsu et al. 1983).
The intrusion of tonalite magma was delayed until
after the subhorizontal displacement of the crust
which occurred subsequent to peak metamorphism.
The time lag corresponds to the time needed for the
magma to be segregated in deep levels and to be
squeezed and rise as a pluton to higher levels. Ac-
cording to Owada et al. (1991), the whole rock
isochron age of the basal tonalites is 56 Ma, while
the mineral isochron age is 52 Ma; the former corre-
sponds to the age of generation, and the latter to
that of consolidation, namely emplacement of the
magma, respectively. The subhorizontal top-to-the-
south displacement of the Hidaka crustal rocks,
therefore, occurred in the time interval between 56
and 52 Ma.
The structural and metamorphic analysis of the
exposed crustal section in the Hidaka belt demon-
strates the essential role of underplating and intru-
sion of mafic magmas in the evolution of magmatic
crust involving crustal metamorphism and anatexis.
The tonalitic intrusives were intruded along struc-
tural breaks such as dkcollement and ramp- and
roof-thrust planes of duplex structures of the crustal
rocks and subsequently mylonitized intensively. This implies that the structural break-forming tectonics
(subhorizontal crustal movement) controlled and en-
hanced the rise and intrusion of the anatectic
magma, but that its generation, rise and intrusion as
plutons into upper levels of the crust controlled and
enhanced the detachment of the upper crust with
subhorizontal movement and subsequent exhuma-
tion of the detached crustal rocks (Toyoshima 1991;
Shimura 1992).
The history of deformation, metamorphism and
magmatism shows that D,, D, and D, stages were
tectonically different. Crustal metamorphism at D,
resulted from underplating and intrusion of mafic
magma. The top-to-the-south subhorizontal displace-
ment of detached lower to upper sequences of crus-
tal rocks at D, in a flat-lying structural state re-
sulted from transcurrent tectonics. Foliations
formed during the D, to D, events were primarily
subhorizontal, but they were inclined to the vertical
through tilting on N-S axes during the D, event.
Retrograde metamorphism during D, shows that the
lower the primary crustal level, the lower the meta-
morphic grade. Because this movement involves
dextral shearing, and the metamorphic rocks were
thrust during the time of dextral shearing, the D,
event resulted from transpressive tectonics
(Fig. 13). This three stage history is the highlight in
the evolution of the Hidaka metamorphic belt.
A /3 /, n 02 tonalite-quartz
Shear zone diorite
Fig. 13 Schematic diagram illustrating the
tectonic movement of the Hidaka metamor-
phic belt (HMB) which occurred after the
intrusion of the D, gabbroic to dioritic rocks,
modified from Toyoshima (1991).
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Tectonic evolution of Hidaka lower crust 197
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S. Wallis for their critical reading and substantial
improvement of the manuscript. Part of this study
was financially supported by a Grant of Fukada
Geological Institute (T.T.) and a Grant-in-Aid of
the Ministry of Education, Science and Culture of
Japan (No. 05402021, M.K.).
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14401738, 1994, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1994.tb00106.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
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Toyoshima (1994) - Tectonic evolution of lower crustal rocks in an exposed magmatic arc section in.txt
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Across-arc Variations in Geochemistry of
Oligocene to Quaternary Basalts from the NE
Japan Arc: Constraints on Source Composition,Mantle Melting and Slab Input Composition
K. Shuto1*, R. Nohara-Imanaka1, M. Sato2, T. Takahashi3, E. Takazawa3,
H. Kawabata4, K. Takanashi5, M. Ban6, N. Watanabe7and
N. Fujibayashi8
1Institute of Science and Technology, Niigata University, 2-8050 Ikarashi, Nishi-ku, Niigata 950-2181, Japan;
2Nissaku Co., Ltd., 1-10-1 Kamikido, Higashi-ku, Niigata 950-0891, Japan;3Department of Geology, Faculty of
Science, Niigata University, 2-8050 Ikarashi, Nishi-ku, Niigata 950-2181, Japan;4Research and Education Faculty,
Kochi University, 2-5-1 Akebono-cho, Kochi 780-8520, Japan;5Geochemical Energy Research and Development
Co., Ltd., 1-22-4, Shinkawa, Chuo-ku, Tokyo 104-0033, Japan;6Earth and Environmental Sciences, Faculty of
Science, Yamagata University, 1-4-12 Kojirakawa-Machi, Yamagata 990-8560, Japan;7Research Institute for
Natural Hazards and Disaster Recovery, Niigata University, 2-8050 Ikarashi, Nishi-ku, Niigata 950-2181, Japan and
8Faculty of Education, Niigata University, 2-8050 Ikarashi, Nishi-ku, Niigata 950-2181, Japan
*Corresponding author. Telephone: þ81-25-261-1297. E-mail: km-shuto@khaki.plala.or.jp
Received April 13, 2015; Accepted November 19, 2015
ABSTRACT
To investigate the nature and origin of across-arc geochemical variations over time in mantle
wedge derived magmas, we have carried out a geochemical study of basalts in the NE Japan arc
spanning an age range from 35 Ma to the present. Back-arc basalts erupted at 24–18 Ma, 10–8 Ma,
6–3 Ma and 2 /C15–0 Ma have higher concentrations of both high field strength elements (HFSE) and
rare earth elements (REE) [particularly light REE (LREE) and middle REE (MREE)], and higher incom-patible trace element ratios compared with frontal-arc basalts at any given time. Geochemical mod-eling of Nb/Yb versus Nb shows that the frontal-arc and back-arc compositional differences areindependent of subduction modification and can, in many cases, be explained by different degreesof melting (higher degrees of melting for frontal-arc magmas and lower degrees of melting for
back-arc magmas) of a nearly homogeneous depleted mid-ocean ridge basalt (MORB) mantle
(DMM)-like source, although there are several exceptions. These include some Pliocene frontal-arcbasalts that may originate from a source that is slightly more depleted than DMM, several35–32 Ma and 24–18 Ma back-arc basalts derived from a lithospheric mantle source that is enrichedin HFSE compared with DMM, and a rare 16–12 Ma basalt that was erupted in the back-arc but wasproduced by a similar degree of melting to frontal-arc basalts erupted at the same time. Variations
in ratios of fluid-mobile and -immobile elements and those of melt-mobile and -immobile elements
for the 35–0 Ma NE Japan basalts indicate that the principal subduction component added to thesource mantle prior to generation of these basalt magmas is a sediment-derived melt. Comparisonof Sr and Nd isotopic compositions for Pacific Ocean MORB, the NE Japan basalts and subductingsediments suggests that the isotopic compositions of most post-16 Ma more depleted back-arc bas-alts can be explained by the addition of <2% bulk sediment; the most enriched isotope compos-
itions of the subcontinental lithosphere-derived magmas can be accounted for by addition of a
maximum 5–7% Japan Trench Sediment (JTS), if the original Sr and Nd compositions of the litho-sphere approximated that of DMM. The Sr and Nd isotope composition of the frontal-arc basaltscan be accounted for by the addition of 1–5% JTS. A depleted asthenospheric mantle (DMM-like)
VCThe Author 2015. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: journals.permissions@oup.com 2257JOURNAL OF
PETROLOGYJournal of Petrology , 2015, Vol. 56, No. 11, 2257–2294
doi: 10.1093/petrology/egv073
Original ArticleDownloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
upwelling model with interaction between asthenospheric mantle-derived magmas and overlying
lithospheric mantle can account for the geochemical characteristics of the 35–0 Ma NE Japanbasalts. The frontal-arc magmas were generally generated by higher degrees of melting of theshallower part of the asthenospheric mantle, whereas the back-arc magmas resulted from lowerdegrees of melting of the deeper part of asthenospheric mantle. These latter magmas underwentinteraction with the lithospheric mantle, resulting in more enriched Sr and Nd isotopic signatures
for the pre-18 Ma back-arc basalts and post-22 Ma frontal-arc basalts, but less interaction, resulting
in more depleted Sr and Nd isotopic signatures, for most of the back-arc basalts younger than16 Ma.
Key words : NE Japan arc; across-arc variation; arc basalt; HFSE; subduction component; DMM;
asthenospheric upwelling; Sr and Nd isotopes
INTRODUCTION
Subduction of the oceanic crust into the mantle and
associated magmatism has played an important role inchemical recycling on Earth. Across-arc variations inthe chemistry of mafic volcanic rocks from both islandarcs and active continental margins can provide import-ant information on the composition of the mantlewedge, and to what degree this has been modified by
addition of a subduction component. Previous studies
of island arcs such as the Aleutians, Kamchatka, NEJapan, Izu–Bonin, the Marianas, Tonga, New Britainand Sangihe, and of active continental margins such asthe Chilean volcanic zone, have revealed that both fluidand melt released from the subducting slab at relatively
low temperature and pressure and high temperature
and pressure, respectively, are transferred into theoverlying mantle wedge (e.g. Elliott et al. , 1997 ;Turner
et al. , 1997 ;Woodhead et al. , 1998 ;Churikova et al. ,
2001 ;Hochstaedter et al. , 2001 ;Ishizuka et al. , 2003 ;
Plank, 2005 ;Kimura & Yoshida, 2006 ;Plank et al. , 2007 ;
Straub et al. , 2010 ,2015 ;Tollstrup et al. , 2010 ;Hanyu
et al. , 2012 ;Jacques et al. , 2013 ,2014 ;Tamura et al. ,
2014 ). These conclusions have been based mainly on
detailed studies of the chemistry of high field strengthelements (HFSE; Ti, Zr, Hf, Nb, Y and Ta), large ion litho-phile elements (LILE; K, Ba, Rb, Sr, Th, U and Pb), rareearth elements (REE), and Sr, Nd, Pb and Hf isotopes.
As the composition of subduction-derived compo-
nents is significantly different from that of the overlyingmantle wedge (e.g. Plank & Langmuir, 1998 ;Taylor &
Nesbitt, 1998 ), addition of a slab component results in
mantle wedge heterogeneity, which has been used toexplain the geochemical variability of arc magmas (e.g.
Leeman et al. , 1994 ;Hochstaedter et al. , 2001 ;Jacques
et al. , 2014 ). Alternatively, it is possible that the sub-arc
mantle is sufficiently heterogeneous, even without theaddition of a subduction component (e.g. Hochstaedter
et al. , 2000 ;Caulfield et al. , 2008 ), to explain the geo-
chemical variations.
In the NE Japan arc, the older (the oldest subducted
crust is early Cretaceous; e.g. Muller et al. , 2008 ) and
cooler crust of the Pacific Plate is subducted beneath
the Eurasian Plate. The arc is bounded to the east by theJapan Trench and to the west by the Japan and Yamato
basins ( Fig. 1 ). Paleomagnetic studies along the NE
Japan arc have suggested that the Japan Sea back-arc
basin formed from 21 to 14 Ma ( Otofuji et al. , 1985 ). The
NE Japan arc records nearly continuous magmatismfrom the early Oligocene to the Quaternary (about35 Ma to the present), spanning the pre-Japan Seaopening to post-opening stages. Thus, NE Japan is an
ideal location to investigate changes in the mantle
wedge and related magma compositions during aperiod of back-arc spreading and how magma compos-ition changes from the volcanic front to the back-arcover more than 30 Myr.
The compositional variations of Quaternary volcanic
rocks from the NE Japan arc are well known in terms of
their major and trace element and radiogenic isotope
compositions. One of the conspicuous characteristics ofthese volcanic rocks is that the concentrations of incom-patible elements including large ion lithophile elements(LILE), some high field strength elements (HFSE) andlight rare earth elements (LREE) increase away from the
volcanic front ( Kuno, 1960 ;Kawano et al. , 1961 ;
Fujimaki & Kurasawa, 1980 ;Sakuyama & Nesbitt, 1986 ;
Nakagawa et al. , 1988 ;Kimura & Yoshida, 2006 ). The
across-arc and along-arc variations in Sr, Nd and Pb iso-topic compositions of Cenozoic NE Japan volcanicrocks have been discussed by several researchers(Notsu, 1983 ;Nohda et al. , 1988 ;Tatsumi et al. , 1988 ;
Togashi et al. , 1992 ;Shuto et al. , 1993 ,2006 ,2013 ;Ujike
& Tsuchiya, 1993 ;Ohki et al. , 1994 ;Kersting et al. , 1996 ;
Gust et al. , 1997 ;Shibata & Nakamura, 1997 ;Kimura &
Yoshida, 2006 ;Yamamoto et al. , 2008 ). However, de-
tailed geochemical across-arc variations for pre-Quaternary volcanic rocks remain unclear, mainly
owing to the lack of comprehensive trace element
chemistry (particularly HFSE and REE). Here, we reportinductively coupled plasma mass spectrometry (ICP-MS) trace element and Sr- and Nd-isotope analyses ofearly Oligocene to Quaternary basaltic rocks from theNE Japan arc in addition to major element analyses. Weuse these new data together with those that we have re-
ported previously to demonstrate the following: (1) at
particular times, NE Japan arc basalts show distinct2258 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Fig. 1. Map showing the location of the study area and the locations of the basalt samples from the NE Japan arc. Those from the
frontal-arc side and back-arc side of the arc are shown as circles and squares, respectively. Numbers within parentheses show theages of the analyzed basalts, based on either K–Ar dating or stratigraphy. Age data sources are described in the text. The distribu-tion of Oligocene to Quaternary volcanic products is simplified from Ono et al. (1981) . The dashed line shows the boundary be-
tween the frontal-arc and back-arc regions.Journal of Petrology , 2015, Vol. 56, No. 11 2259Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
across-arc geochemical variations in terms of the abun-
dances and ratios of incompatible trace elements, al-though these are less clear for the period 16–12 Ma; (2)sediment-derived melt is the major subduction compo-nent added to the mantle source; (3) arc basalt compos-
itions reflect variable degrees of melting of source
mantle that includes a depleted mid-ocean ridge basalt(MORB) mantle (DMM)-like source, a source slightlymore depleted than DMM and subcontinental litho-sphere to which has been added a small amount ( <5%)
of sediment melt.
GEOLOGICAL, PETROLOGICAL AND TECTONIC
BACKGROUND
Recently, Yoshida et al. (2014) reviewed late Cenozoic
magmatism in the NE Japan arc and examined the rela-tionship between the magmatism and the crust–mantle
structure. They divided the Cenozoic magmatism into
three periods: continental margin (66–21 Ma), back-arc(21–13 /C15 Ma) and island-arc (13 /C15–0 Ma). Magmatism
during each period is linked with geodynamic changesin the mantle, associated crustal heating and remelting,and subsequent crustal cooling and strong compres-sion. Here, we focus on the relationship between the
petrology of the early Oligocene to Quaternary volcanic
rocks of the NE Japan arc, and their relationship to tec-tonism. Throughout this period, the NE Japan arc expe-rienced almost continuous volcanism over a distance ofabout 100–150 km extending from central Honshu toSW Hokkaido ( Fig. 1 ). Based on radiometric age data for
Cenozoic volcanic rocks from the NE Japan arc, most
basalts were erupted in the early Oligocene (35–32 Ma),early Miocene (24–18 Ma), middle Miocene (16–12 Ma),late Miocene (10–8 Ma), Pliocene (6–3 Ma) andQuaternary (2 /C15 Ma to the present).
At the commencement of Japan Sea opening, early
Oligocene basaltic volcanism related to back-arc rifting
took place largely on land in a limited area of the pre-
sent-day back-arc margin of the NE Japan arc ( Fig. 1 ).
After the rifting stage, the volcanic front, which waslocated along the western Japan Sea coast of the pre-sent NE Japan arc, moved trenchward, and volcanic ac-tivity is recorded along the entire NE Japan arc during
the early to middle Miocene. As a result, the position of
the 16–12 Ma volcanic front is located 30–50 km east ofthe Quaternary volcanic front ( Fig. 1 ;Ohki et al. , 1993 ;
Shuto, 2009 ). The eastern limit of volcanic rocks
younger than middle late Miocene (about 8 Ma) almostagrees with the position of the Quaternary volcanicfront ( Tamura & Shuto, 1989 ;Shuto & Yashima, 1990 ).
This indicates that the 16–12 Ma volcanic front may
have started to move westward at about 12 Ma, and hasbeen fixed in its present position since about 8 Ma ( Ohki
et al. , 1993 ).
The early Oligocene (35–32 Ma) volcanic rocks found
on Okushiri Island and the Oga Peninsula are mainly
calc-alkaline andesite and dacite, and normal tholeiitic
basalt and high-TiO
2basalt characterized by moderateto high concentrations of TiO 2and other HFSE
(Yamamoto et al. , 1991 ;Okamura et al. , 1993 ;Fukase &
Shuto, 2000 ). Adakitic andesite of the Matsue Basalt
Formation, dated at 34 /C14 Ma, has been found from the
southern part of Okushiri Island ( Sato et al. , 2014 ).
Small amounts of 33–29 Ma andesite and more felsic
volcanic rocks are also located about 50 km behind theQuaternary volcanic front in the central part of the NEJapan arc ( Shuto, 2009 ). From this, the 35–32 Ma
Matsue basalt on southern Okushiri Island and theKamo basalt on the Oga Peninsula ( Fig. 1 ) are classified
as back-arc basalts in this study.
The widespread early Miocene (24–18 Ma) volcanic
rocks largely comprise both tholeiitic and calc-alkalinebasalts through to rhyolite. High-TiO
2basalt has also
been observed on the frontal-arc side as well as on theback-arc side. Also, exposures of adakitic andesite havebeen reported from the 23–20 Ma Aonaegawa
Formation on southern Okushiri Island ( Sato et al. ,
2013 ). Basaltic and associated intermediate to felsic vol-
canic rocks erupted in the back-arc at 35–20 Ma are geo-chemically analogous to those from continental riftzones, such as the Rio Grande Rift of New Mexico andColorado (e.g. Barberi et al. , 1982 ), which led several
workers to argue that these back-arc volcanic rocks
were formed in association with rifting of the Eurasiancontinental margin arc during the pre-opening stage ofthe Japan Sea ( Fukase & Shuto, 2000 ;Shuto et al. ,
2006 ).
It is estimated that the total volume of middle
Miocene (16–12 Ma) volcanic rocks erupted in the NE
Japan arc is much greater than that associated with the
Pliocene to Quaternary volcanism ( Tatsumi et al. , 1989 ).
In particular, 16–12 Ma back-arc volcanism was domi-nated by bimodal volcanism within submarine grabens,which produced voluminous basaltic and felsic (daciticto rhyolitic) eruptive products ( Konda, 1974 ;Yamaji &
Sato, 1989 ;Shuto et al. , 1993 ). It must be emphasized
that the 16–12 Ma volcanic rocks from the NE Japan arcshow no across-arc variation in total alkali and K
2O con-
tents (e.g. Shuto et al. , 2013 ,fig. 2 ), in contrast to the
Quaternary volcanic rocks.
Several workers have argued that the geochemical
characteristics of 16–12 Ma basalts from the NE Japan
arc and back-arc submarine volcanism may reflect up-
welling of depleted hot asthenosphere into the subcon-tinental lithosphere beneath the back-arc margin of thearc, coincident with the spreading of the Japan Seaback-arc basin (e.g. Nohda et al. , 1988 ;Tatsumi et al. ,
1989 ;Shuto et al. , 1993 ,2006 ;Ohki et al. , 1994 ;Yagi
et al. , 2001 ). Such hot asthenospheric flow would mod-
ify the thermal structure of the mantle wedge beneaththe frontal arc during the early to middle Miocene, andresult in partial melting of both mantle wedge peridotiteand the cool subducting Pacific Plate ( Yamamoto &
Hoang, 2009 ;Shuto et al. , 2013 ).
Volcanism during the earliest late Miocene (10–8 Ma)
to Pliocene (6–3 Ma) in the NE Japan arc was less in-
tense. The volcanic rocks emplaced at this time show a2260 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Fig. 2. SiO 2vs FeO*/MgO, SiO 2vs Na 2OþK2O, SiO 2vs K 2O, SiO 2vs TiO 2and SiO 2vs P 2O5relations for NEJ basalts. (a) Early
Oligocene (35–32 Ma), (b) early Miocene (24–18 Ma), (c) middle Miocene (16–12 Ma), (d) late Miocene (10–8 Ma), (e) Pliocene (6–
3 Ma), and (f) Quaternary (2 /C15–0 Ma). The fields of TH (tholeiitic series) and CA (calc-alkaline series) for SiO 2–(FeO*/MgO) are from
Miyashiro (1974) ; the sub-alkalic and alkalic fields for SiO 2–(Na 2OþK2O) are from Miyashiro (1978) , and the LK (low-K andesite),
MK (medium-K andesite) and HK (high-K andesite) fields for SiO 2–K2O are from Gill (1981) .Journal of Petrology , 2015, Vol. 56, No. 11 2261Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
relatively wide compositional range, including both tho-
leiitic and calc-alkaline basalt to dacite ( Tamura &
Shuto, 1989 ;Shuto & Yashima, 1990 ). Quaternary
(2/C15–0 Ma) volcanoes in NE Japan belong to either the
frontal chain (Nasu volcanic zone) or the back-arc chain(Chokai volcanic zone) ( Kawano et al. , 1961 ). Systematic
differences between them in terms of the petrographyand geochemistry of the volcanic rocks have been re-
ported by Kuno (1966) ,Sakuyama & Nesbitt (1986) ,
Nakagawa et al. (1988) andKimura & Yoshida (2006) .Fig. 2. Continued.2262 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
SAMPLES
Basaltic rocks from both the NE Japan frontal-arc and
back-arc regions (NEJ frontal-arc basalts and NEJ back-arc basalts) were collected from 17 and 28 locations, re-
spectively ( Fig. 1 ). In this paper, we define the region
including the Quaternary Chokai volcanic zone itselfand the more western zone as the back-arc, whereasthe frontal arc is defined by the region to the east of theChokai volcanic zone. The boundary between them isshown by the bold dashed line in Fig. 1 .
We have already carried out petrographical examin-
ation and major and/or trace element analysis of basalts
from most of the volcanic fields shown in Fig. 1 (Shuto
et al. ,1 9 8 5 ,2006 ,2013 ;Tamura & Shuto, 1989 ;Shuto &
Yashima, 1990 ;Asakusa Collaborative Research Group,
1991 ;Yamamoto et al. ,1 9 9 1 ;Ohki et al. ,1 9 9 4 ,1995 ;
Takimoto & Shuto, 1994 ;Nakajima et al. ,1 9 9 5 ;Banet al. ,
1997 ;Fukase & Shuto, 2000 ;Kondo et al. ,2 0 0 0 ;Sato
et al. ,2 0 0 7 ,2013 ,2014 ;Watanabe et al. ,2 0 0 9 ;Takanashi
et al. ,2 0 1 1 ). Petrographic and geochemical data for bas-
alts from the Inaniwadake (Pliocene), Takadate (earlyMiocene) and Tenmyosan (early Miocene) districts andthe Quaternary Iwate, Funagata, Kampu, Sannome-gataand Chokai volcanoes have been reported by other re-
searchers (e.g. Wada, 1981 ;Hayashi, 1984 ;Maruyama
et al. ,1 9 8 8 ;Uto et al. ,1 9 8 9 ;Yoshinaga & Nakagawa,
1999 ;Yasui & Yamamoto, 2006 ;Yamamoto & Hoang,
2009 ;Kuritani et al. ,2 0 1 4 a,2014 b
).
Because pre-Quaternary basalts from the NE Japan
arc have undergone some degree of post-magmatic al-
teration (particularly the middle Miocene submarine
volcanic rocks found in the back-arc side), we have care-fully collected the freshest samples from the respectivelocations. Thus, plagioclase, clinopyroxene and ortho-pyroxene phenocrysts in the basalts discussed here aregenerally fresh, although olivine phenocrysts are partlyor wholly altered to clay minerals. Eruption ages are
based in large part on radiometric age data, but also on
stratigraphic relations that have been reported by Sato
et al. (2007) ,Shuto (2009) andTakanashi et al. (2011) for
most of the basalts and associated volcanic rocks fromeach location discussed here. Eruption ages for otherbasalt samples are assumed to be 20 Ma for the
Tenmyosan district ( Yamamoto & Hoang, 2009 ), 5/C16–
5/C12 Ma for the Pliocene Kurohanayama district ( Ban
et al. , 1997 ), 0/C11 Ma for Iwate volcano ( Ito & Doi, 2005 ),
0/C19 Ma for Funagata volcano ( Umeda et al. , 1999 ),
0/C102 Ma for Sannome-gata volcano ( Kitamura, 1990 )
and middle Pleistocene for Asakusa volcano ( Asakusa
Collaborative Research Group, 1991 ).
ANALYTICAL METHODS
Whole-rock major and trace element analyses
All samples were crushed to 3–5 mm diameter frag-
ments, which were then washed, dried and pulverized
in an agate ball mill. Whole-rock major element com-
positions were determined by X-ray fluorescence (XRF)spectrometry, using a Rigaku Rix 3000 system at the
Faculty of Science, Niigata University. Glass disks wereprepared by fusing 0 /C15 g of rock powder with lithium tet-
raborate (Li
2B4O7,5/C10 g). Prior to fusion, the sample
powders were heated at 880/C14C for 2 /C15 h to measure loss
on ignition. Analytical procedures have been described
byTakahashi & Shuto (1997) .
The concentrations of trace elements were analyzed
by ICP-MS (Agilent 7500a) at the Faculty of Science,Niigata University. To avoid incomplete dissolution of zir-con and other refractory minerals, we prepared samplesolutions using a combined Na
2CO3fusion and acid di-
gestion procedure. Powdered sample material (0 /C11g )
was fused with Na 2CO3(0/C15 g) at 1050/C14C. The solution
was neutralized with HNO 3þHCl and diluted by a factor
of 20 000 with a mixture of HF–HNO 3–HCl. Average val-
ues of trace element data for USGS reference materialBHVO-2, determined by repeated analyses ( n¼11) in our
laboratory using an HF–HNO
3acid digestion method, are
consistent with BHVO-1 values of Eggins et al. (1997)
within 2% deviation. Thus we used a single solution ofBHVO-2 as an external calibration standard, using the ref-erence BHVO-1 values of Eggins et al. (1997) , except for
Tm, for which we used the Tm value of BHVO-1 reported
byMakishima & Nakamura (2006) . Variation during the
analytical runs was corrected for using three internalstandards (In, Re, Bi). External standardization was per-formed for single elements in unknown samples by inter-polation of results for replicate analyses of BHVO-2 after5–6 unknown samples. Analytical quality was determinedby comparing our results for USGS reference material
JB-2 with literature values ( Supplementary Data
Electronic Appendix Table A1 ;supplementary data are
available for downloading at http://www.petrology.oxfordjournals.org). Our data are consistent with literature val-ues within 9% for most elements.
Sr and Nd isotope analyses
Mass spectrometric measurements of each sample forSr and Nd isotopes were conducted using a MAT262mass spectrometer at the Faculty of Science, NiigataUniversity, following the procedures of Miyazaki &
Shuto (1998) . Extraction of Sr and Nd from rock powder
has been described by Takahashi et al. (2009) and
Kagami et al. (1987) , respectively. Measured
87Sr/86Sr
and143Nd/144Nd ratios were normalized to
86Sr/88Sr¼0/C11194 and146Nd/144Nd¼0/C17219, respect-
ively. The87Sr/86Sr composition of NIST987 was meas-
ured during this study, averaging 0 /C1710251 60/C1000003
(2r,n¼20);143Nd/144Nd ratios for JNdi-1 gave a mean
value of 0 /C1512106 60/C1000003 (2 r,n¼17).
GEOCHEMICAL ACROSS-ARC VARIATIONS OF
POST-35 Ma TO THE PRESENT BASALTS FROM
THE NE JAPAN ARC
We provide new major and trace element analyses
for 82 samples from 27 locations in the NE Japan arcJournal of Petrology , 2015, Vol. 56, No. 11 2263Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
(Fig. 1 ). Of these samples, 47 were analysed for Sr and
Nd isotopes. All the data are listed in Tables 1 and 2, re-
spectively. In addition to the new data we also includeanalyses reported elsewhere ( Shuto et al. , 2006 ,2013 ;
Sato et al. , 2007 ,2013 ,2014 ;Takanashi et al. , 2011 ) for
47 basaltic samples from 18 locations ( Supplementary
Data Electronic Appendix Table A2 ).
All the major element analyses have been recalcu-
lated to 100 wt % on an anhydrous basis. A total of 129samples discussed here have MgO >5 wt %, and
SiO
2<54/C16 wt %, which we refer to as ‘basalts’ for
simplicity.
Major element compositions
Because the pre-Quaternary NEJ basalts have beensubjected to post-magmatic alteration, original mag-matic abundances of LILE such as K and Na may have
been modified. However, such alteration effects appear
minimal, as all of these basalts plot in the unalteredfield in the Na
2OþK2O versus Na 2O/K 2O plot proposed
byMiyashiro (1975) . In addition, the Sr content of bas-
alts is easily changed as a result of alteration, althoughNd is very resistant to alteration. However, Sr and Ndisotopic compositions are similar for an altered dolerite
from North Hokkaido and fresh clinopyroxene sepa-
rated from the same rock, which contains chlorite andcalcite ( Shuto et al. , 2004 ). Petrographic observation
shows that the degree of alteration is smaller in most ofthe basalt samples analyzed in this study than in thisdolerite. Thus, we assume that alteration did not signifi-
cantly affect the Sr and Nd isotopic compositions and
major and trace element compositions of the studiedbasalts.
Most of the NEJ frontal-arc and back-arc basalts, irre-
spective of their ages, show spatial distributions strad-dling the discrimination line between the tholeiitic andcalc-alkaline suites of Miyashiro (1974) in terms of
SiO
2–FeO*/MgO ( Fig. 2 a–f). The 24–18 Ma, 10–8 Ma,
6–3 Ma and 2 /C15–0 Ma NEJ basalts exhibit increase in
Na2OþK2O and K 2O contents from the frontal arc to the
back-arc ( Fig. 2 b,d ,e and f). More exactly, the NEJ fron-
tal-arc basalts have Na 2OþK2O plotting in the lower
part of the sub-alkalic field of Miyashiro (1978) in terms
of SiO 2–total alkalis, whereas the data for the back-arc
basalts plot in the upper part of the same field.Similarly, the frontal-arc basalts have lower K
2O, falling
mostly in the Low-K field and partly around the Low-Kto Medium-K boundary of Gill (1981) , compared with
the back-arc basalts, which plot mostly in the Medium-Kfield and partly in the upper part of the Low-K field. The
Na
2OþK2O values of the 35–32 Ma back-arc basalts are
also higher than those of the 24–18 Ma, 10–8 Ma, 6–3 Maand 2 /C15–0 Ma frontal-arc basalts, but are similar to those
of the 24–18 Ma back-arc basalts ( Fig. 2 a,b ,d , e and f).
Also, K
2O contents of the 35–32 Ma back-arc basalts are
similar to those of the 24–18 Ma back-arc basalts. In
contrast, the Na 2OþK2O contents of both the 16–12 Ma
frontal-arc and back-arc basalts overlap in thesub-alkalic field on the SiO 2–total alkalis diagram, and
K2O contents overlap in the SiO 2–K2O diagram ( Fig. 2 c).
Some of the 35–12 Ma and 24–18 Ma back-arc basalts(the Matsue, Aonaegawa and Upper Fukuyama basalts)have high TiO
2(/C242 wt %) and P 2O5contents ( Fig. 2 a and
b). All other 35–0 Ma back-arc basalts have TiO 2
contents within a relatively narrow range from 0 /C18t o
1/C12 wt %. The 24–18 Ma and 10–8 Ma frontal-arc basalts
are slightly lower in TiO 2than the back-arc basalts
erupted at the same time, whereas TiO 2contents of
both the frontal-arc and back-arc basalts at 6–3 Ma and2/C15–0 Ma overlap, and those at 16–12 Ma partly overlap
in terms of SiO
2–TiO 2. On an SiO 2–P2O5diagram, the
P2O5contents of the 24–18 Ma, 10–8 Ma, 6–3 Ma, and
2/C15–0 Ma back-arc basalts are higher than those of fron-
tal-arc basalts, although the 16–12 Ma frontal- and back-arc basalts show some overlap.
MgO versus incompatible element
concentrations and ratios
Niobium, Ta, Zr, Hf, Y, Yb and Lu are insensitive to the
subduction component whether it is a fluid or a sedi-
ment melt ( Pearce & Peate, 1995 ). Therefore, the rela-
tionships between both MgO and the concentrations ofHFSE such as Nb and Zr, and MgO and trace elementratios such as Nb/Yb, Zr/Y, Zr/Lu and La/Yb for the35 Ma to present NEJ basalts ( Figs 3 –7) can provide fun-
damental information about the nature. of the mantle
wedge, independent of subduction modification andmagmatic processes within the wedge.
The 24–18 Ma frontal-arc basalts have lower Nb and
Zr contents at a given MgO than similar age back-arcbasalts ( Fig. 3 ). Similar across-arc variations in these
elements are also observed in the 10–8 Ma, 6–3 Ma and
2/C15–0 Ma basalts ( Figs 5 –7). Moreover, frontal-arc bas-
alts have lower Nb/Yb, Zr/Y, Zr/Lu and La/Yb at thesame MgO compared with back-arc basalts at any time(Figs 3 ,5,6and 7). However, such differences in elem-
ent contents or ratios are not apparent for 16–12 Mafrontal-arc and back-arc basalts at MgO contents lower
than about 8 wt % ( Fig. 4 ), and back-arc basalts erupted
at 35–32 Ma have similar incompatible element con-tents and ratios to 24–18 Ma back-arc basalt magmas(Fig. 3 ).
The 2 /C15–0 Ma back-arc basalts have higher Th/Nb
and Ba/Nb at a given MgO content compared with fron-tal-arc basalts of the same age ( Figs 7 g and h). In con-
trast, there is no similar systematic variation with MgO
content between the frontal-arc and back-arc basaltsproduced at 24–18 Ma, 16–12 Ma, 10–8 Ma and 6–3 Ma(Figs 3 g, h, 4g, h, 5 g, h, and 6 g, h).
Age versus incompatible element concentrations
and ratios
We plot incompatible element concentrations and ratios
in the studied basalts (17 frontal-arc basalts and 28
back-arc basalts) versus their formation ages ( Fig. 8 ). In
the case of basalts for which we have two or more2264 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
analyzed samples, we used their average concentra-
tions and ratios. In Fig. 8 , the incompatible element
ratios of DMM, normal mid-ocean ridge basalt
(N-MORB), enriched mid-ocean ridge basalt (E-MORB)and ocean island basalt (OIB) are also shown for com-parison. The key aspects illustrated by Fig. 8 are as fol-
lows: (1) most of the frontal-arc basalts at 24–18 Ma,10–8 Ma, 6–3 Ma and 2 /C15–0 Ma, characterized by lower
Nb and Zr contents, and lower Nb/Yb, Zr/Y, Zr/Lu and
La/Yb ratios compared with the back-arc basalts for the
same period, show a relatively narrow range in incom-patible element ratios from levels corresponding tothose of DMM to those of E-MORB, whereas the back-arc basalts exhibit a wide range from values corres-ponding to those of N-MORB to those of OIB; (2)
amongst the back-arc basalts, the 35–32 Ma and
24–18 Ma basalts generally have higher incompatibleelement ratios than the 10–8 Ma, 6–3 Ma and 2 /C15–0 Ma
basalts. Two Pliocene Inaniwadake and Araya basalts
from the frontal arc have lower Nb/Yb, Zr/Y and Zr/Lu
ratios than DMM, and thus are more depleted com-pared with all other frontal-arc basalts; (3) both the fron-tal- and back-arc basalts with ages of 16–12 Ma cannotbe discriminated from each other in terms of incompat-ible element concentrations and ratios, and their incom-patible element ratios are mostly higher than those of
N-MORB, but markedly lower than those of OIB, al-
though some frontal-arc basalts have Zr/Y and Zr/Luratios lying between DMM and N-MORB ratios. Wefurther examine these features observed in 16–12 Mabasalts using age versus Nb/Y and Zr/Y ( Fig. 9 ) plots in
which single analyses of basalts are shown. The frontal-
arc Ryozen, Rokugo Nakuidake, and Tomari basalts
have roughly similar Nb/Yb ratios to the frontal-arcTable 1 : Major- and trace-element compositions of the NEJ basalts determined by XRF and ICP-MS, respectively
Age: Early Oligocene (35–32 Ma) Early Miocene (24–18 Ma)
Region: Back-arc Frontal arc
Basalt name: Kamo Takadate Tenmyosan
Sample no.: 084 127 KH07 KH16 TA03 TA05 TA06 TA10 TM01 TM02 TM10 TM11 TM12
Age (Ma): 32 32 32 32 22 22 22 22 20 20 20 20 20
SiO 2(wt %) 50 /C157 51 /C167 51 /C154 51 /C132 51 /C178 49 /C174 51 /C160 51 /C150 48 /C186 48 /C129 49 /C182 50 /C113 49 /C195
TiO 2 1/C141 1 /C114 1 /C149 1 /C115 0 /C187 0 /C180 0 /C186 0 /C185 0 /C168 0 /C167 0 /C173 0 /C172 0 /C172
Al2O3 17/C125 18 /C172 17 /C145 17 /C105 18 /C119 16 /C169 18 /C116 18 /C115 14 /C186 14 /C171 16 /C115 16 /C114 16 /C115
FeO* 8 /C131 7 /C137 7 /C170 7 /C168 8 /C179 9 /C115 8 /C170 8 /C159 9 /C135 9 /C139 8 /C178 8 /C169 8 /C169
MnO 0 /C116 0 /C116 0 /C114 0 /C119 0 /C118 0 /C118 0 /C118 0 /C119 0 /C117 0 /C118 0 /C116 0 /C116 0 /C116
MgO 6 /C103 5 /C105 5 /C143 6 /C157 5 /C186 7 /C153 5 /C195 5 /C182 11 /C177 11 /C169 9 /C128 9 /C170 9 /C130
CaO 7 /C190 8 /C178 7 /C173 9 /C125 9 /C162 10 /C120 9 /C144 9 /C150 10 /C136 10 /C155 10 /C144 10 /C131 10 /C137
Na2O3 /C152 3 /C183 3 /C180 3 /C110 2 /C167 2 /C124 2 /C169 2 /C166 2 /C114 1 /C192 2 /C139 2 /C141 2 /C140
K2O0 /C198 0 /C181 1 /C100 0 /C181 0 /C135 0 /C130 0 /C146 0 /C147 0 /C124 0 /C125 0 /C130 0 /C132 0 /C130
P2O5 0/C135 0 /C121 0 /C134 0 /C124 0 /C115 0 /C110 0 /C115 0 /C114 0 /C107 0 /C107 0 /C108 0 /C109 0 /C108
LOI 3 /C117 2 /C106 3 /C119 2 /C138 1 /C158 2 /C139 1 /C144 1 /C143 1 /C167 2 /C120 1 /C161 1 /C143 1 /C138
Total 99 /C164 99 /C180 99 /C181 99 /C175 100 /C103 99 /C131 99 /C162 99 /C131 100 /C116 99 /C193 99 /C175 100 /C110 99 /C150
Sc (ppm) 13 /C101 2 /C141 1 /C161 4 /C182 8 /C113 3 /C182 8 /C142 8 /C113 6 /C173 6 /C143 5 /C143 1 /C182 8 /C13
V 116 119 100 99 /C17 224 254 224 223 238 229 255 240 249
Rb 6 /C195 5 /C116 7 /C107 7 /C196 4 /C156 3 /C198 9 /C118 11 /C103 /C196 3 /C159 5 /C129 5 /C193 5 /C195
Sr 619 618 500 613 329 248 337 339 183 182 246 243 248Y2 2 /C181 7 /C132 2 /C151 8 /C191 9 /C141 9 /C131 9 /C162 0 /C121 7 /C111 6 /C171 7 /C141 7 /C121 7 /C12
Zr 160 92 /C17 160 103 57 /C124 2 /C195 8 /C135 7 /C103 7 /C103 6 /C154 0 /C164 1 /C184 1 /C16
Nb 8 /C171 4 /C199 9 /C182 5 /C128 3 /C100 1 /C181 3 /C102 2 /C1
95 0 /C1823 0 /C1771 1 /C134 1 /C147 1 /C140
Cs 0 /C1406 0 /C1066 3 /C163 0 /C1789 0 /C1381 0 /C1160 0 /C1265 0 /C1321 0 /C1053 0 /C1046 0 /C1069 0 /C1079 0 /C1097
Ba 219 228 226 192 158 121 170 166 66 /C176 5 /C129 7 /C109 9 /C129 7 /C12
La 19 /C191 0 /C151 5 /C141 3 /C175 /C157 4 /C101 5 /C150 5 /C156 2 /C141 2 /C136 3 /C140 3 /C157 3 /C144
Ce 45 /C192 4 /C173 5 /C193 2 /C141 3 /C171 0 /C121 4 /C111 3 /C166 /C153 6 /C144 8 /C141 8 /C172 8 /C151
Pr 6 /C114 3 /C141 4 /C180 4 /C146 1 /C193 1 /C148 1 /C195 1 /C190 1 /C105 1 /C104 1 /C128 1 /C130 1 /C127
Nd 25 /C131 4 /C182 0 /C151 9 /C149 /C134 7 /C121 9 /C107 9 /C121 5 /C151 5 /C138 6 /C100 6 /C127 6 /C123
Sm 5 /C122 3 /C136 4 /C157 4 /C119 2 /C136 2 /C108 2 /C142 2 /C120 1 /C164 1 /C173 1 /C182 1 /C188 1 /C181
Eu 1 /C158 1 /C119 1 /C156 1 /C136 0 /C1981 0 /C1812 0 /C1938 0 /C1948 0 /C1701 0 /C1662 0 /C1774 0 /C1781 0 /C1734
Gd 4 /C194 3 /C150 4 /C163 3 /C199 3 /C110 2 /C177 3 /C108 3 /C107 2 /C125 2 /C114 2 /C136 2 /C136 2 /C125
Tb 0 /C1751 0 /C1572 0 /C1718 0 /C1627 0 /C1493 0 /C1474 0 /C1511 0 /C1480 0 /C1418 0 /C1422 0 /C1403 0 /C1407 0 /C1408
Dy 4 /C140 3 /C135 4 /C145 3 /C166 2 /C198 3 /C105 2 /C194 3 /C116 2 /C158 2 /C155 2 /C171 2 /C170 2 /C164
Ho 0 /C1921 0 /C1700 0 /C1881 0 /C1758 0 /C1679 0 /C1691 0 /C1662 0 /C1683 0 /C1579 0 /C1568 0 /C1591 0 /C1602 0 /C1582
Er 2 /C152 1 /C198 2 /C157 2 /C119 1 /C196 1 /C195 2 /C107 2 /C100 1 /C175 1 /C174 1 /C175 1 /C169 1 /C164
Tm 0 /C1360 0 /C1286 0 /C1363 0 /C1324 0 /C1307 0 /C1287 0 /C1280 0 /C1293 0 /C1266 0 /C1245 0 /C1274 0 /C1261 0 /C1245
Yb 2 /C133 1 /C189 2 /C150 1 /C199 1 /C181 1 /C177 1 /C173 1 /C179 1 /C174 1 /C181 1 /C179 1 /C186 1 /C180
Lu 0 /C1331 0 /C1273 0 /C1339 0 /C1289 0 /C1301 0 /C1288 0 /C1303 0 /C1291 0 /C1263 0 /C1248 0 /C1279 0 /C1271 0 /C1281
Hf 3 /C176 2 /C153 3 /C183 2 /C170 1 /C158 1 /C125 1 /C155 1 /C149 1 /C105 1 /C103 1 /C115 1 /C115 1 /C118
Ta 0 /C1531 0 /C1370 0 /C1721 0 /C1361 0 /C1168 0 /C1111 0 /C1192 0 /C1188 0 /C1049 0 /C1051 0 /C1072 0 /C1079 0 /C1091
Pb 5 /C199 6 /C132 5 /C120 5 /C175 4 /C116 3 /C120 4 /C118 4 /C107 1 /C162 1 /C145 2 /C122 2 /C137 2 /C119
Th 1 /C171 1 /C123 1 /C104 1 /C122 0 /C1632 0 /C1561 0 /C1718 0 /C1682 0 /C1379 0 /C1372 0 /C1490 0 /C1533 0 /C1442
U0 /C1438 0 /C1302 0 /C1340 0 /C1350 0 /C1173 0 /C1126 0 /C1179 0 /C1182 0 /C1090 0 /C1076 0 /C1128 0 /C1139 0 /C1143
(continued)Journal of Petrology , 2015, Vol. 56, No. 11 2265Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
basalts produced after 12 Ma ( Fig. 9 a), and the former
three basalts also have similar Zr/Y ratios to these fron-tal-arc basalts ( Fig. 9 b). However, the remaining frontal-
arc basalts (Ishimoriyama and Washinosu basalts) havehigh Nb/Yb and Zr/Yb ratios similar to those of the back-arc basalts younger than 12 Ma. These basalts(Ishimoriyama basalt IS03 and Washinosu basalt
0628E) are differentiated, and characterized by enriched
Sr and Nd isotope compositions compared with otherfrontal-arc 16–12 Ma basalts, and thus their elevatedNb/Yb and Zr/Yb may reflect crustal assimilation and/orcrystal fractionation. The high Zr/Y ratios of the Tomaribasalts may also be affected by crustal assimilation be-
cause of their unusually high Th/Nb and Ba/Nb ratios
(Fig. 4 ). Therefore, most of the relatively unfractionatedand uncontaminated 16–12 Ma frontal-arc basalts (i.e.
Ryozen, Rokugo and Nakuidake basalts), as well as theless differentiated Ishimoriyama and Washinotsu bas-alts, have similar or slightly higher Nb/Yb and Zr/Ybratios compared with frontal-arc basalts younger than12 Ma ( Fig. 9 a and b). The 13 /C13 Ma back-arc Tappizaki
basalt has Nb/Yb and Zr/Yb ratios lower than those of
other back-arc basalts of a similar age, but similar to
those of the frontal-arc Ryozen, Rokugo and Nakuidakebasalts. Thus, excluding localized basalts (e.g. Tappizakibasalt), similar to the 24–18 Ma, 10–8 Ma, 6–3 Ma and2/C15–0 Ma basalts, those erupted at 16–12 Ma in both the
frontal-arc and back-arc regions are generally character-
ized by low Nb/Yb and Zr/Yb ratios, and high Nb/Yb and
Zr/Yb ratios, respectively.Table 1 : Continued
Age: Early Miocene (24–18 Ma) Middle Miocene (16–12 Ma)
Region: Frontal arc Frontal arc Back-arc
Basalt name: Tenmyosan Rokugo Nakui- Washinosu Tappizaki L. Tobi Ogi
dake shima
Sample no.: TM13 TM14 TM20 MG15 0604 WA06 0628E TP23A KS03 TP14 TO13 SW01N SW01S
Age (Ma): 20 20 20 15 15 12 12 13 12 /C151 4 /C14 1 51 21 2
SiO 2(wt %) 48 /C182 49 /C131 50 /C107 50 /C124 50 /C152 47 /C180 50 /C152 49 /C137 51 /C172 50 /C177 50 /C100 48 /C104 48 /C135
TiO 2 0/C168 0 /C168 0 /C172 1 /C119 0 /C160 0 /C175 0 /C181 0 /C186 0 /C179 1 /C127 0 /C191 0 /C189 0 /C160
Al2O3 15/C136 15 /C138 15 /C148 16 /C111 18 /C112 19 /C161 18 /C193 17 /C187 19 /C167 16 /C166 17 /C103 15 /C162 15 /C162
FeO* 9 /C116 9 /C101 8 /C129 11 /C101 9 /C189 8 /C114 7 /C174 9 /C134 8 /C132 11 /C184 8 /C152 7 /C163 7 /C100
MnO 0 /C118 0 /C117 0 /C115 0 /C117 0 /C117 0 /C109 0 /C109 0 /C121 0 /C116 0 /C113 0 /C120 0 /C112 0 /C112
MgO 10 /C174 10 /C188 10 /C122 5 /C162 5 /C156 4 /C186 5 /C128 7 /C125 5 /C110 5 /C143 6 /C175 7 /C198 8 /C172
CaO 10 /C153 10 /C153 9 /C132 9 /C123 10 /C184 9 /C168 9 /C195 10 /C166 10 /C170 9 /C158 11 /C142 12 /C101 13 /C158
Na2O2 /C113 2 /C119 2 /C122 2 /C119 1 /C171 2 /C123 2 /C162 1 /C194 2 /C151 2 /C111 2 /C124 2 /C108 1 /C152
K2O0 /C123 0 /C120 0 /C141 0 /C121 0 /C126 0 /C141 0 /C143 0 /C131 0 /C124 0 /C152 0 /C116 0 /C181 0 /C147
P2O5 0/C107 0 /C107 0 /C109 0 /C117 0 /C106 0 /C112 0 /C112 0 /C114 0 /C112 0 /C114 0 /C112 0 /C118 0 /C113
LOI 1 /C162 1 /C136 2 /C146 4 /C108 2 /C101 6 /C110 3 /C136 1 /C125 0 /C190 0 /C185 2 /C140 4 /C166 3 /C186
Total 99 /C154 99 /C179 99 /C143 100 /C123 99 /C173 99 /C181 99 /C185 99 /C120 100 /C123 99 /C130 99 /C174 100 /C103 99 /C197
Sc (ppm) 34 /C123 5 /C193 0 /C193 6 /C184 7 /C113 3 /C162 8 /C173 6 /C192 9 /C144 4 /C183 8 /C133 5 /C194 1 /C14
V 244 241 234 349 424 272 216 283 246 512 245 261 243
Rb 2 /C162 3 /C184 10 /C112 /C173 4 /C129 4 /C117 13 /C123 /C161 2 /C175 6 /C104 2 /C181 11 /C125 /C124
Sr 206 210 179 255 184 327 363 331 406 287 309 329 322Y1 7 /C181 7 /C171 7 /C192 2 /C111 5 /C191 6 /C172 0 /C112 0 /C122 0 /C132 1 /C172 1 /C112 0 /C131 5 /C13
Zr 37 /C153 7 /C155 4 /C114 4 /C152 5 /C125 0 /C136 6 /C105 5 /C175 5 /C115 0 /C167 2 /C116 7 /C104 0 /C10
Nb 1 /C132 1 /C136 2 /C150 1 /C161 2 /C125 2 /C1
21 3 /C184 3 /C106 2 /C169 1 /C188 3 /C104 8 /C156 2 /C140
Cs 0 /C1071 0 /C1183 0 /C1246 0 /C1287 0 /C1608 0 /C1064 0 /C1834 0 /C1244 0 /C1543 0 /C1113 0 /C1748 0 /C1394 0 /C1037
Ba 75 /C157 6 /C15 103 74 /C15 131 128 124 81 /C17 103 174 62 /C14 107 48 /C13
La 3 /C112 3 /C118 4 /C170 4 /C109 4 /C133 6 /C100 6 /C185 7 /C133 6 /C129 5 /C163 4 /C173 7 /C171 4 /C148
Ce 7 /C172 8 /C101 11 /C101 1 /C128 /C183 14 /C101 6 /C171 5 /C141 4 /C181 3 /C171 2 /C131 7 /C101 0 /C16
Pr 1 /C118 1 /C118 1 /C153 1 /C163 1 /C112 1 /C184 2 /C129 2 /C106 1 /C192 1 /C193 1 /C178 2 /C127 1 /C161
Nd 5 /C162 5 /C185 6 /C195 8 /C186 5 /C106 8 /C136 10 /C159 /C122 9 /C146 8 /C198 8 /C144 9 /C187 7 /C133
Sm 1 /C179 1 /C178 2 /C102 2 /C158 1 /C140 2 /C120 2 /C169 2 /C136 2 /C154 2 /C167 2 /C144 2 /C135 1 /C188
Eu 0 /C1722 0 /C1710 0 /C1748 1 /C107 0 /C1501 0 /C1843 0 /C1948 0 /C1932 0 /C1927 0 /C1971 0 /C1918 0 /C1962 0 /C1770
Gd 2 /C127 2 /C132 2 /C131 3 /C145 1 /C198 2 /C165 3 /C105 2 /C195 2 /C190 3 /C121 3 /C126 3 /C100 2 /C127
Tb 0 /C1410 0 /C1412 0 /C1429 0 /C1561 0 /C1342 0 /C1433 0 /C1518 0 /C1524 0 /C1504 0 /C1577 0 /C1510 0 /C1522 0 /C1412
Dy 2 /C173 2 /C164 2 /C164 3 /C154 2 /C137 2 /C171 3 /C108 3 /C118 3 /C134 3 /C156 3 /C129 3 /C120 2 /C142
Ho 0 /C1603 0 /C1621 0 /C1604 0 /C1790 0 /C1538 0 /C1591 0 /C1700 0 /C1762 0 /C1694 0 /C1813 0 /C1753 0 /C1709 0 /C1546
Er 1 /C180 1 /C183 1 /C172 2 /C125 1 /C171 1 /C170 1 /C199 2 /C107 2 /C100 2 /C129 2 /C121 1 /C197 1 /C154
Tm 0 /C1272 0 /C1267 0 /C1265 0 /C1318 0 /C1245 0 /C1231 0 /C1283 0 /C1336 0 /C1304 0 /C1315 0 /C1318 0 /C1315 0 /C1232
Yb 1 /C180 1 /C174 1 /C179 2 /C105 1 /C152 1 /C151 1 /C165 2 /C102 1 /C180 2 /C103 1 /C196 2 /C104 1 /C151
Lu 0 /C1271 0 /C1292 0 /C1270 0 /C1312 0 /C1228 0 /C1231 0 /C1283 0 /C1310 0 /C1309 0 /C1332 0 /C1321 0 /C1290 0 /C1205
Hf 1 /C110 1 /C115 1 /C141 1 /C124 0 /C179 1 /C138 1 /C168 1 /C158 1 /C146 1 /C148 1 /C181 1 /C152 0 /C194
Ta 0 /C1081 0 /C1082 0 /C1150 0 /C1091 0 /C1141 0 /C1128 0 /C1203 0 /C1198 0 /C1172 0 /C1111 0 /C1151 0 /C1550 0 /C1181
Pb 1 /C198 2 /C171 2 /C147 1 /C188 3 /C151 3 /C109 3 /C156 3 /C125 4 /C195 3 /C160 2 /C144 1 /C109 0 /C196
Th 0 /C1491 0 /C1558 0 /C1800 0 /C1501 0 /C1920 1 /C126 1 /C112 1 /C120 1 /C119 1 /C119 0 /C1352 0 /C1938 0 /C1451
U0 /C1137 0 /C1156 0 /C1168 0 /C1161 0 /C1249 0 /C1778 0 /C1282 0 /C1372 0 /C1284 0 /C1339 0 /C1077 0 /C1547 0 /C1216
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It is also clear that Nb/Yb, Zr/Y, and Zr/Lu for back-arc
basalts gradually decrease from 35 Ma to the present.The 35–32 Ma and 24–18 Ma basalts with higher ratiosalso have moderate to high contents of K
2O,
K2OþNa2O, TiO 2and P 2O5, and thus these basalts are
considered to be analogous to those found in continen-tal rift zones, such as the Rio Grande Rift ( Shuto et al. ,
2006 ;Sato et al. , 2007 ). These within-plate type basalts
have OIB-like Nb/Yb, Zr/Y, and Zr/Lu ratios ( Fig. 8 ).
N-MORB normalized incompatible element
patterns
N-MORB normalized multi-element plots for the studied
NEJ basalts are shown in Fig. 10 . Both frontal-arc andback-arc basalts erupted at 24–18 Ma, 16–12 Ma, 10–8 Ma,
6–3 Ma and 2 /C15–0 Ma exhibit steeply inclined patterns
with enrichment in LILE and LREE, whereas HFSE suchas Nb, Ta, Zr and Hf are significantly depleted comparedwith the adjacent LILE and LREE ( Fig. 10 b–f). These pat-
terns are typical of island arc basalts reported elsewhere(e.g. Pearce et al. , 2005 ). Apart from the 16–12 Ma basalts,
frontal arc basalts are more depleted in Nb, Ta, Zr and Hf
compared with back-arc basalts at any particular time(Fig. 10 b and d–f). The 16–12 Ma basalts also have arc-
like trace element patterns, but the patterns are similarfor both the front- and back-arc ( Fig. 10 c). We also note
that the 35–32 Ma back-arc basalts have similar patterns
and degrees of depletion in Nb, Ta, Zr and Hf to 24–18 Ma
back-arc basalts ( Fig. 10 a and b).Table 1 : Continued
Age: Middle Late Miocene (10–8 Ma)
Miocene
(16–12 Ma)
Region: Back-arc Frontal arc Back-arc
Basalt name: Ogi Joge Mitaki Aosawa Aizome
Sample no.: SR01 71 SI14 J18 TA1 TA2 TA3 NM03 NM24 NM42 AZ01 AZ02 AZ03
Age (Ma): 12 8 /C128 /C128 /C127 /C197 /C197 /C191 0 /C141 0 /C141 0 /C141 0 /C121 0 /C121 0 /C12
SiO 2(wt %) 49 /C151 50 /C198 50 /C100 52 /C114 50 /C124 49 /C197 50 /C109 48 /C103 47 /C160 47 /C163 50 /C105 49 /C191 50 /C125
TiO 2 1/C103 0 /C154 0 /C162 0 /C193 0 /C168 0 /C169 0 /C171 0 /C198 1 /C113 1 /C103 0 /C189 0 /C188 0 /C190
Al2O3 15/C124 17 /C144 17 /C189 16 /C107 17 /C150 17 /C163 17 /C146 17 /C173 16 /C120 16 /C172 16 /C184 17 /C113 17 /C114
FeO* 7 /C196 8 /C159 9 /C113 10 /C187 9 /C178 9 /C166 9 /C168 7 /C118 7 /C179 8 /C159 8 /C107 7 /C180 7 /C171
MnO 0 /C115 0 /C117 0 /C120 0 /C121 0 /C118 0 /C118 0 /C118 0 /C122 0 /C118 0 /C118 0 /C112 0 /C107 0 /C107
MgO 8 /C101 5 /C191 5 /C183 6 /C103 7 /C105 6 /C163 6 /C169 6 /C155 7 /C117 7 /C139 6 /C185 6 /C193 6 /C178
CaO 10 /C171 10 /C181 10 /C147 10 /C161 11 /C142 11 /C139 11 /C138 9 /C110 10 /C133 9 /C187 9 /C102 9 /C105 8 /C195
Na2O2 /C109 1 /C160 1 /C152 2 /C101 1 /C164 1 /C163 1 /C164 2 /C163 2 /C153 2 /C143 2 /C159 2 /C138 2 /C135
K2O0 /C155 0 /C113 0 /C109 0 /C118 0 /C115 0 /C115 0 /C115 0 /C153 0 /C126 0 /C144 0 /C178 0 /C157 0 /C157
P2O5 0/C120 0 /C105 0 /C105 0 /C108 0 /C105 0 /C105 0 /C106 0 /C122 0 /C121 0 /C121 0 /C121 0 /C118 0 /C118
LOI 4 /C122 3 /C106 3 /C194 0 /C194 1 /C154 2 /C117 2 /C110 6 /C165 6 /C155 5 /C150 4 /C149 5 /C109 4 /C194
Total 99 /C167 99 /C129 99 /C174 100 /C107 100 /C122 100 /C114 100 /C114 99 /C182 99 /C195 99 /C198 99 /C191 99 /C197 99 /C183
Sc (ppm) 34 /C124 1 /C134 0 /C184 8 /C105 9 /C126 1 /C176 2 /C135 1 /C194 7 /C114 9 /C144 9 /C145 9 /C166 2 /C11
V 267 266 298 345 395 397 414 346 369 387 365 415 435Rb 7 /C190 1 /C173 1 /C122 2 /C168 4 /C157 4 /C172 3 /C182 10 /C152 /C159 9 /C198 21 /C171 4 /C111 1 /C13
Sr 260 174 174 187 222 224 223 432 319 326 378 325 327
Y2 3 /C121 4 /C161 4 /C172 1 /C101 9 /C151 9 /C102 0 /C122 6 /C162 9 /C122 9 /C102 7 /C122 5 /C112 6 /C19
Zr 84 /C112 8 /C142 3 /C183 6 /C123 1 /C143 1 /C143 3 /C119 1 /C189 8 /C139 2 /C16 101 /C138 4 /C167 0 /C14
Nb 9 /C142 1 /C111 0 /C184 1 /C128 1 /C107 1 /C1
08 1 /C113 4 /C116 4 /C181 4 /C191 4 /C153 3 /C164 3 /C183
Cs 0 /C1263 0 /C1062 0 /C1066 0 /C1098 0 /C1167 0 /C1124 0 /C1083 0 /C1085 0 /C1010 0 /C1075 0 /C1674 0 /C1346 0 /C1565
Ba 91 /C185 4 /C123 5 /C116 3 /C105 8 /C196 0 /C186 2 /C11 122 95 /C19 129 209 181 193
La 9 /C136 2 /C142 1 /C186 2 /C168 2 /C145 2 /C149 2 /C163 8 /C137 8 /C157 8 /C104 10 /C186 9 /C112 9 /C143
Ce 20 /C176 /C107 4 /C184 7 /C117 6 /C134 6 /C137 6 /C175 21 /C112 1 /C122 0 /C132 5 /C152 1 /C162 1 /C19
Pr 2 /C176 0 /C188 0 /C174 1 /C113 1 /C102 1 /C100 1 /C107 3 /C105 3 /C111 2 /C196 3 /C147 2 /C196 3 /C106
Nd 12 /C114 /C140 3 /C184 5 /C174 5 /C104 5 /C106 5 /C115 13 /C181 4 /C111 3 /C121 4 /C151 2 /C191 3 /C11
Sm 2 /C177 1 /C123 1 /C138 1 /C186 1 /C173 1 /C179 1 /C174 3 /C161 3 /C178 3 /C162 3 /C167 3 /C137 3 /C138
Eu 1 /C108 0 /C1561 0 /C1520 0 /C1763 0 /C1678 0 /C1692 0 /C1751 1 /C134 1 /C132 1 /C128 1 /C130 1 /C119 1 /C120
Gd 3 /C164 1 /C178 1 /C185 2 /C169 2 /C118 2 /C120 2 /C131 3 /C197 4 /C115 3 /C189 3 /C187 3 /C167 3 /C178
Tb 0 /C1641 0 /C1343 0 /C1318 0 /C1472 0 /C1427 0 /C1420 0 /C1460 0 /C1671 0 /C1701 0 /C1693 0 /C1678 0 /C1621 0 /C1661
Dy 3 /C168 2 /C122 2 /C113 3 /C106 2 /C185 2 /C192 2 /C199 4 /C130 4 /C165 4 /C139 4 /C113 3 /C194 4 /C111
Ho 0 /C1840 0 /C1542 0 /C1523 0 /C1747 0 /C1628 0 /C1621 0 /C1640 0 /C1849 0 /C1957 0 /C1890 0 /C1851 0 /C1841 0 /C1863
Er 2 /C127 1 /C155 1 /C152 2 /C121 1 /C180 1 /C183 1 /C192 2 /C147 2 /C172 2 /C172 2 /C146 2 /C138 2 /C156
Tm 0 /C1358 0 /C1245 0 /C1231 0 /C1317 0 /C1286 0 /C1291 0 /C1310 0 /C1371 0 /C1425 0 /C1412 0 /C1399 0 /C1364 0 /C1388
Yb 2 /C130 1 /C162 1 /C158 2 /C105 1 /C189 1 /C175 1 /C190 2 /C134 2 /C158 2 /C161 2 /C141 2 /C142 2 /C142
Lu 0 /C1333 0 /C1250 0 /C1243 0 /C1311 0 /C1287 0 /C1302 0 /C1316 0 /C1367 0 /C1430 0 /C1400 0 /C1401 0 /C1394 0 /C1418
Hf 1 /C191 0 /C190 0 /C170 1 /C111 0 /C190 0 /C190 0 /C191 2 /C106 2 /C113 1 /C199 2 /C113 1 /C187 1 /C188
Ta 0 /C1620 0 /C1057 0 /C1048 0 /C1087 0 /C1080 0 /C1077 0 /C1069 0 /C1265 0 /C1292 0 /C1292 0 /C1292 0 /C1220 0 /C1208
Pb 1 /C166 1 /C169 1 /C151 1 /C192 2 /C110 0 /C198 0 /C193 0 /C182 0 /C148 1 /C126 3 /C122 3 /C135 3 /C126
Th 1 /C135 0 /C1577 0 /C1315 0 /C1409 0 /C1400 0 /C1396 0 /C1434 1 /C116 1 /C109 1 /C110 1 /C193 1 /C156 1 /C152
U0 /C1395 0 /C1142 0 /C1083 0 /C1095 0 /C1086 0 /C1088 0 /C1096 0 /C1330 0 /C1290 0 /C1300 0 /C1504 0 /C1377 0 /C1342
(continued)Journal of Petrology , 2015, Vol. 56, No. 11 2267Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Rare earth element chemistry
Chondrite-normalized REE patterns for NEJ basalts
show a distinct difference between frontal-arc and back-arc basalts erupted at 24–18 Ma, 10–8 Ma, 6–3 Ma and
2/C15–0 Ma ( Fig. 11 b and d–f). For basalts with similar
MgO contents, most back-arc basalts have higher REEcontents compared with those found in the frontal arc,particularly in terms of the LREE (La to Nd) and middleREE (MREE; Sm to Dy), which results in steeper REEpatterns in back-arc basalts. This may mean that the dif-
ferent REE patterns cannot be ascribed to different de-
grees of fractional crystallization between the frontalarc and back-arc, but may reflect fundamental differ-ences in REE (especially LREE and MREE) contents be-tween frontal-arc and back-arc primary basalt magmas.Similar to incompatible element plots ( Fig. 10 c), frontal-
arc and back-arc basalts erupted at 16–12 Ma have over-
lapping REE patterns over a similar MgO range (about9–5 wt %), although the frontal-arc basalts withMgO>10 wt % have flatter patterns compared with
MgO-poor frontal-arc and back-arc basalts ( Fig. 11 c).
The REE patterns of 35–32 Ma back-arc basalts aresimilar to those of the 24–18 Ma back-arc basalts with
similar MgO contents ( Fig. 11 a and b).
In summary, the 24–18 Ma, 10–8 Ma, 6–3 Ma and
2/C15–0 Ma frontal-arc and back-arc basalts can be sepa-
rated in terms of incompatible element contents andratios. In general, back-arc basalts have higher REE con-tents and steeper REE patterns than the frontal-arc bas-
alts. Only back-arc basalts were erupted at 35–32 Ma,Table 1 : Continued
Age: Late Miocene (10–8 Ma) Pliocene (6–3 Ma)
Region: Back-arc Frontal arc
Basalt name: Aizome Upper Tobishima Kurohanayama Araya Hakaseyama
Sample no.: AZ04 TB14 TB15 TB25a TB27 090901 090902 100803 AR01 AR02 90403 401 504
Age (Ma): 10 /C129 /C139 /C139 /C139 /C135 /C165 /C165 /C16 332 /C182 /C182 /C18
SiO 2(wt %) 50 /C121 51 /C184 50 /C121 51 /C189 51 /C150 51 /C133 51 /C195 51 /C156 50 /C168 50 /C172 49 /C168 51 /C196 50 /C189
TiO 2 0/C182 1 /C114 1 /C130 1 /C123 1 /C113 0 /C164 0 /C173 0 /C183 0 /C180 0 /C179 0 /C177 0 /C175 0 /C176
Al2O3 15/C167 16 /C122 19 /C143 15 /C173 16 /C132 17 /C109 17 /C123 16 /C113 16 /C176 16 /C168 17 /C136 17 /C121 17 /C139
FeO* 7 /C175 8 /C137 6 /C154 8 /C170 8 /C136 9 /C158 9 /C197 10 /C167 10 /C138 10 /C109 9 /C143 8 /C177 9 /C124
MnO 0 /C114 0 /C124 0 /C109 0 /C121 0 /C128 0 /C118 0 /C118 0 /C120 0 /C119 0 /C118 0 /C119 0 /C117 0 /C118
MgO 7 /C160 6 /C178 7 /C189 7 /C138 6 /C183 7 /C103 6 /C157 6 /C131 7 /C141 6 /C161 6 /C173 5 /C182 6 /C170
CaO 9 /C151 9 /C180 8 /C114 8 /C171 9 /C187 10 /C193 10 /C156 10 /C181 11 /C120 11 /C132 11 /C125 10 /C174 10 /C175
Na2O2 /C125 2 /C181 3 /C108 2 /C189 2 /C174 1 /C157 1 /C161 1 /C179 1 /C159 1 /C157 2 /C113 2 /C130 2 /C115
K2O0 /C158 0 /C140 0 /C126 0 /C147 0 /C134 0 /C113 0 /C116 0 /C114 0 /C110 0 /C112 0 /C144 0 /C175 0 /C155
P2O5 0/C116 0 /C122 0 /C125 0 /C127 0 /C121 0 /C105 0 /C105 0 /C107 0 /C106 0 /C106 0 /C115 0 /C114 0 /C114
LOI 5 /C106 2 /C115 2 /C150 2 /C110 2 /C141 1 /C152 1 /C173 1 /C100 1 /C101 1 /C189 1 /C131 0 /C194 1 /C120
Total 99 /C176 99 /C197 99 /C169 99 /C159 100 /C100 100 /C104 100 /C174 99 /C151 100 /C117 100 /C101 99 /C144 99 /C156 99 /C196
Sc (ppm) 52 /C105 6 /C106 3 /C195 2 /C195 7 /C104 5 /C104 5 /C144 8 /C167 3 /C107 4 /C106 7 /C183 4 /C113 4 /C15
V 362 378 411 370 375 286 303 336 509 507 430 248 254
Rb 8 /C178 21 /C151 /C101 38 /C191 7 /C191 /C163 2 /C156 1 /C184 1 /C105 1 /C194 7 /C120 16 /C101 0 /C19
Sr 382 336 395 411 336 155 159 178 247 251 421 297 321Y2 8 /C113 7 /C113 7 /C144 1 /C163 6 /C121 5 /C191 9 /C121 8 /C152 3 /C142 4 /C172 6 /C142 1 /C122 0 /C10
Zr 88 /C12 108 125 146 109 26 /C163 1 /C143 1 /C172 4 /C152 4 /C165 1 /C165 1 /C144 7 /C19
Nb 5 /C111 6 /C151 6 /C119 10 /C104 /C198 0 /C1842 1 /C106 1 /C114 0 /C1752 0 /C1
791 1 /C172 2 /C101 1 /C160
Cs 0 /C1599 3 /C104 0 /C1195 10 /C163 /C147 0 /C1089 0 /C1133 0 /C1048 0 /C1039 0 /C1103 0 /C1506 0 /C1405 0 /C1725
Ba 324 201 125 256 196 54 /C166 4 /C145 4 /C126 7 /C117 1 /C10 284 253 223
La 12 /C131 1 /C101 1 /C161 5 /C131 0 /C172 /C132 2 /C190 2 /C146 2 /C116 2 /C113 8 /C191 8 /C164 7 /C109
Ce 27 /C152 6 /C132 8 /C123 5 /C172 5 /C155 /C130 6 /C151 6 /C148 5 /C122 5 /C114 18 /C191 7 /C181 5 /C15
Pr 3 /C159 3 /C175 4 /C107 4 /C189 3 /C160 0 /C184 1 /C106 0 /C198 0 /C199 1 /C102 2 /C171 2 /C154 2 /C115
Nd 14 /C181 6 /C161 7 /C192 1 /C111 6 /C124 /C134 5 /C114 5 /C119 5 /C144 5 /C169 11 /C181 1 /C119 /C192
Sm 3 /C155 4 /C154 4 /C175 5 /C142 4 /C108 1 /C135 1 /C160 1 /C168 1 /C195 2 /C101 3 /C120 2 /C183 2 /C156
Eu 1 /C120 1 /C152 1 /C166 1 /C167 1 /C144 0 /C1562 0 /C1701 0 /C1702 0 /C1738 0 /C1801 1 /C112 0 /C1921 0 /C1896
Gd 3 /C177 5 /C117 5 /C138 5 /C170 4 /C183 1 /C195 2 /C126 2 /C130 2 /C169 2 /C190 3 /C137 3 /C110 3 /C100
Tb 0 /C1630 0 /C1862 0 /C1931 0 /C1992 0 /C1801 0 /C1348 0 /C1408 0 /C1416 0 /C1491 0 /C1521 0 /C1611 0 /C1510 0 /C1478
Dy 4 /C110 5 /C145 5 /C188 6 /C120 5 /C135 2 /C128 2 /C170 2 /C173 3 /C131 3 /C147 3 /C186 3 /C111 3 /C110
Ho 0 /C1870 1 /C115 1 /C121 1 /C126 1 /C109 0 /C1543 0 /C1632 0 /C1631 0 /C1710 0 /C1738 0 /C1801 0 /C1722 0 /C1693
Er 2 /C152 3 /C123 3 /C146 3 /C168 3 /C109 1 /C168 1 /C199 1 /C196 2 /C112 2 /C118 2 /C124 2 /C112 1 /C196
Tm 0 /C1407 0 /C1501 0 /C1550 0 /C1595 0 /C1494 0 /C1244 0 /C1280 0 /C1280 0 /C1313 0 /C1348 0 /C1369 0 /C1308 0 /C1327
Yb 2 /C151 3 /C116 3 /C118 3 /C153 3 /C110 1 /C159 1 /C177 1 /C182 1 /C195 2 /C120 2 /C140 2 /C117 1 /C191
Lu 0 /C1401 0 /C1503 0 /C1538 0 /C1579 0 /C1487 0 /C1251 0 /C1272 0 /C1283 0 /C1349 0 /C1357 0 /C1371 0 /C1300 0 /C1312
Hf 2 /C119 3 /C104 3 /C130 3 /C183 2 /C190 0 /C184 0 /C190 0 /C190 0 /C169 0 /C175 1 /C143 1 /C134 1 /C135
Ta 0 /C1371 0 /C1390 0 /C1365 0 /C1630 0 /C1345 0 /C1057 0 /C1053 0 /C1070 0 /C1042 0 /C1049 0 /C1126 0 /C1107 0 /C1102
Pb 4 /C155 4 /C160 4 /C139 5 /C143 4 /C133 1 /C163 2 /C114 1 /C149 1 /C184 2 /C101 4 /C166 4 /C149 5 /C191
Th 2 /C187 1 /C102 1 /C105 1 /C156 0 /C1968 0 /C1459 0 /C1590 0 /C1424 0 /C1325 0 /C1341 2 /C107 2 /C153 1 /C192
U0 /C1734 0 /C1240 0 /C1298 0 /C1331 0 /C1205 0 /C1081 0 /C1111 0 /C1089 0 /C1082 0 /C1075 0 /C1477 0 /C1616 0 /C1521
(continued)2268 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
and these have similar HFSE abundances, incompatible
element ratios and REE chemistry to 24–18 Ma back-arcbasalts. These across-arc variations are less distinct infrontal-arc and back-arc basalts erupted at 16–12 Ma.
Sr and Nd isotope chemistry
We calculated initial Sr and Nd isotope ratios (Sr Iand
NdI,Table 2 ) based on the respective eruption ages for
selected basalts from the study area. All Sr Iand Nd Ival-
ues of frontal-arc basalts, regardless of eruption age(22–0 /C11 Ma), are undepleted compared with N-MORB.
They cover a wide range in Sr
Ifrom 0 /C1703921 to
0/C1705047, and Nd Ifrom 0 /C1512661 to 0 /C1512883 ( Table 2 ).
Although the 22 /C14 Ma Lower Fukuyama back-arc basalts
have similar Sr Iand Nd Ivalues (Sr I¼0/C1704388 and0/C1704612, and Nd I¼0/C1512601 and 0 /C1512665) to frontal-
arc basalts, the Sr Ivalues of other back-arc basalts
younger than 10 /C14 Ma range from 0 /C1702970 to 0 /C1703949,
which is mostly less than the Sr Irange of the frontal-arc
basalts. The Nd Ivalues of these back-arc basalts range
from 0 /C1512787 to 0 /C1512956, most of which are higher
than Nd Ivalues of the frontal-arc basalts ( Table 2 ).
DISCUSSION
Possibility of crustal assimilation during basalt
production
Seismic velocity data show that crustal thickness in the
present NE Japan arc varies from about 40 km to 26 km,
and decreases from the land area towards both theTable 1 : Continued
Age: Pliocene (6–3 Ma) Quaternary (2 /C15–0 Ma)
Region: Frontal arc Back-arc Frontal arc
Basalt name: Hakase- Inaniwadake Manaita- Myojin-iwa Ajarayama Funagata
yama yama
Sample no.: 420 IN10 IN11 IN12 IN13 MIY06 L138 L141 4-01A 3-11A FU02 FU05 Fu06
Age (Ma): 2 /C182 /C172 /C172 /C172 /C173 /C16 331 /C191 /C190 /C190 /C190 /C19
SiO 2(wt %) 50 /C197 51 /C169 51 /C151 52 /C100 51 /C192 51 /C106 54 /C124 53 /C139 51 /C175 50 /C186 51 /C142 51 /C143 51 /C151
TiO 2 0/C179 0 /C197 0 /C197 0 /C195 0 /C194 1 /C106 0 /C198 1 /C104 0 /C187 0 /C192 0 /C182 0 /C184 0 /C183
Al2O3 16/C139 15 /C179 15 /C184 16 /C105 16 /C101 17 /C174 17 /C112 17 /C159 17 /C128 17 /C167 17 /C197 18 /C107 18 /C104
FeO* 9 /C164 11 /C152 11 /C152 11 /C132 11 /C131 8 /C140 6 /C159 7 /C166 10 /C100 9 /C171 9 /C133 9 /C139 9 /C133
MnO 0 /C118 0 /C121 0 /C121 0 /C121 0 /C121 0 /C117 0 /C116 0 /C113 0 /C119 0 /C119 0 /C117 0 /C118 0 /C118
MgO 7 /C182 5 /C178 5 /C174 5 /C171 5 /C181 4 /C198 4 /C192 5 /C114 6 /C110 5 /C170 6 /C141 6 /C153 6 /C148
CaO 10 /C114 10 /C161 10 /C135 10 /C155 10 /C138 10 /C110 9 /C139 9 /C183 10 /C169 10 /C194 10 /C154 10 /C139 10 /C149
Na2O2 /C107 1 /C170 1 /C164 1 /C168 1 /C164 2 /C183 2 /C175 2 /C167 2 /C110 2 /C127 2 /C114 2 /C106 2 /C114
K2O0 /C153 0 /C110 0 /C112 0 /C113 0 /C113 0 /C158 1 /C161 1 /C122 0 /C114 0 /C123 0 /C121 0 /C120 0 /C119
P2O5 0/C112 0 /C106 0 /C106 0 /C106 0 /C106 0 /C134 0 /C118 0 /C118 0 /C110 0 /C111 0 /C108 0 /C108 0 /C108
LOI 1 /C172 1 /C127 1 /C160 1 /C135 1 /C159 2 /C163 1 /C145 1 /C141 0 /C189 0 /C199 0 /C196 1 /C126 1 /C102
Total 100 /C137 99 /C170 99 /C156 100 /C100 100 /C100 99 /C189 99 /C139 100 /C126 100 /C112 99 /C159 100 /C106 100 /C142 100 /C129
Sc (ppm) 35 /C165 4 /C155 3 /C135 2 /C135 2 /C103 0 /C193 0 /C113 2 /C174 0 /C173 7 /C123 4 /C173 6 /C103 3 /C15
V 241 411 403 400 400 229 203 255 298 248 265 269 258
Rb 11 /C120 /C156 0 /C188 1 /C135 1 /C132 6 /C189 42 /C163 0 /C150 /C190 4 /C103 3 /C164 2 /C100 1 /C193
Sr 288 167 182 177 177 401 367 394 256 288 269 263 258Y2 0 /C142 0 /C182 1 /C162 1 /C102 1 /C133 1 /C122 8 /C142 3 /C182 4 /C112 3 /C131 7 /C191 7 /C171 7 /C13
Zr 49 /C152 5 /C192 6 /C142 5 /C152 5 /C11 114 106 91 /C153 7 /C194 2 /C184 3 /C194 3 /C154 2 /C11
Nb 1 /C123 0 /C1722 0 /C1713 0 /C1719 0 /C1690 4 /C176 5 /C133 4 /C1
81 1 /C140 1 /C170 1 /C162 1 /C158 1 /C158
Cs 0 /C1295 0 /C1032 0 /C1021 0 /C1034 0 /C1035 0 /C1253 1 /C188 0 /C1413 0 /C1069 0 /C1237 0 /C1271 0 /C1065 0 /C1146
Ba 168 66 /C166 9 /C106 9 /C146 9 /C10 217 327 253 94 /C179 9 /C14 108 106 104
La 6 /C110 1 /C185 2 /C107 2 /C107 2 /C109 13 /C151 2 /C181 0 /C194 /C136 3 /C157 3 /C173 3 /C160 3 /C155
Ce 14 /C175 /C133 5 /C159 5 /C129 5 /C137 33 /C102 9 /C112 5 /C179 /C121 9 /C114 9 /C124 9 /C104 8 /C197
Pr 2 /C100 0 /C189 0 /C195 0 /C194 0 /C196 4 /C149 3 /C184 3 /C140 1 /C170 1 /C153 1 /C138 1 /C135 1 /C134
Nd 9 /C179 4 /C189 5 /C113 5 /C113 5 /C118 20 /C131 7 /C131 5 /C178 /C125 7 /C194 6 /C159 6 /C154 6 /C148
Sm 2 /C145 1 /C176 1 /C173 1 /C179 1 /C168 4 /C186 4 /C116 3 /C167 2 /C134 2 /C135 2 /C107 1 /C194 1 /C184
Eu 0 /C1862 0 /C1701 0 /C1778 0 /C1718 0 /C1743 1 /C164 1 /C125 1 /C116 0 /C1969 1 /C100 0 /C1809 0 /C1802 0 /C1782
Gd 3 /C121 2 /C165 2 /C174 2 /C153 2 /C166 5 /C108 4 /C139 4 /C130 3 /C127 3 /C135 2 /C155 2 /C143 2 /C148
Tb 0 /C1491 0 /C1532 0 /C1518 0 /C1509 0 /C1502 0 /C1803 0 /C1704 0 /C1607 0 /C1527 0 /C1538 0 /C1442 0 /C1441 0 /C1432
Dy 3 /C121 3 /C120 3 /C131 3 /C108 3 /C119 5 /C103 4 /C138 4 /C106 3 /C140 3 /C157 2 /C165 2 /C171 2 /C171
Ho 0 /C1701 0 /C1752 0 /C1787 0 /C1729 0 /C1756 1 /C104 0 /C1962 0 /C1842 0 /C1768 0 /C1791 0 /C1589 0 /C1587 0 /C1601
Er 2 /C110 2 /C100 2 /C111 1 /C186 2 /C102 3 /C110 2 /C181 2 /C137 2 /C119 2 /C119 1 /C178 1 /C171 1 /C169
Tm 0 /C1304 0 /C1353 0 /C1339 0 /C1308 0 /C1334 0 /C1427 0 /C1378 0 /C1353 0 /C1317 0 /C1346 0 /C1274 0 /C1263 0 /C1262
Yb 1 /C191 2 /C122 2 /C119 2 /C128 2 /C125 2 /C164 2 /C145 2 /C102 1 /C199 1 /C199 1 /C173 1 /C172 1 /C157
Lu 0 /C1291 0 /C1302 0 /C1335 0 /C1321 0 /C1309 0 /C1437 0 /C1413 0 /C1310 0 /C1307 0 /C1321 0 /C1263 0 /C1251 0 /C1268
Hf 1 /C141 0 /C188 0 /C180 0 /C182 0 /C183 2 /C180 2 /C183 2 /C143 1 /C112 1 /C122 1 /C119 1 /C115 1 /C118
Ta 0 /C1071 0 /C1041 0 /C1038 0 /C1039 0 /C1048 0 /C1238 0 /C1374 0 /C1338 0 /C1090 0 /C1108 0 /C1090 0 /C1094 0 /C1088
Pb 3 /C144 2 /C152 2 /C139 2 /C144 2 /C134 4 /C165 4 /C199 5 /C187 2 /C184 3 /C112 2 /C127 2 /C143 2 /C121
Th 1 /C138 0 /C1345 0 /C1373 0 /C1348 0 /C1385 1 /C140 3 /C152 2 /C191 0 /C1449 0 /C1477 0 /C1617 0 /C1627 0 /C1593
U0 /C1366 0 /C1065 0 /C1077 0 /C1082 0 /C1081 0 /C1365 1 /C139 0 /C1659 0 /C1121 0 /C1137 0 /C1143 0 /C1136 0 /C1139
(continued)Journal of Petrology , 2015, Vol. 56, No. 11 2269Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Pacific Ocean and the Japan Sea coasts ( Zhao et al. ,
1992 ;Miyamichi et al. , 1994 ). Assuming a similar crustal
thickness over the last 35 Myr, crustal assimilationcould have taken place along the entire NE Japan arc. Ifthis is the case, there should be a clear correlation be-tween the isotopic compositions of crustal rocks andbasalts.
Based on seismological and geological data, the pre-
sent upper crust at depths of less than 15 km in the
back-arc region of the central NE Japan arc consistslargely of pre-Cenozoic sedimentary rocks andCretaceous to Paleogene granitic rocks ( Sato et al. ,
2007 ). These basement rocks have more radiogenic Sr
I
(at 22 Ma) and less radiogenic Nd I(at 22 Ma) than early
Miocene (22–18 Ma) NEJ back-arc basalts ( Sato et al. ,
2007 , and references therein), and so the Sr and Ndisotopic variations of these basalts could result from as-
similation of crustal rocks. The Sr isotope composition
of Quaternary NEJ frontal-arc volcanic rocks and pre-Cenozoic granitic rocks led Kimura & Yoshida (2006) to
argue that the volcanic rocks were highly affected bycrustal assimilation. However, Sato et al. (2007) showed
that along-arc isotopic variations in early Miocene NEJback-arc basalts could not be correlated with isotopic
variations of Cretaceous to Paleogene granitic rocks.
Further, simple mixing trajectories between basalticand granitic compositions in terms of Sr
I–K2O/TiO 2,S r I–
K2O/P 2O5and Sr I–Rb/Zr indicated that crustal assimila-
tion is unlikely in the petrogenesis of early MioceneNEJ back-arc basalts ( Sato et al. , 2007 ).
However, amongst the 35–0 Ma NEJ basalts, there
are several basalts with unusually higher Ba/Zr, Th/Zr,Table 1 : Continued
Age: Quaternary (2 /C15–0 Ma)
Region: Frontal arc Back-arc
Basalt name: Iwate Asakusa Chokai Kampu
Sample no.: IW01 IW02 IW03 IW04 IW05 IW06 14201 35202 54109 61106 CH02 81710 82008
Age (Ma): 0 /C110 /C110 /C110 /C110 /C110 /C11 111100 /C110 /C11
SiO 2(wt %) 51 /C150 51 /C126 51 /C126 51 /C153 51 /C101 51 /C140 49 /C164 49 /C136 47 /C192 47 /C165 53 /C178 52 /C108 52 /C170
TiO 2 0/C178 0 /C178 0 /C177 0 /C178 0 /C178 0 /C178 0 /C193 0 /C195 0 /C199 0 /C197 0 /C192 0 /C183 0 /C180
Al2O3 17/C102 16 /C192 16 /C183 16 /C193 16 /C185 16 /C196 19 /C115 19 /C178 20 /C103 19 /C180 17 /C180 18 /C127 18 /C114
FeO* 9 /C133 9 /C140 9 /C149 9 /C138 9 /C152 9 /C140 9 /C114 9 /C116 9 /C150 9 /C140 7 /C170 7 /C180 7 /C167
MnO 0 /C117 0 /C117 0 /C117 0 /C117 0 /C117 0 /C117 0 /C118 0 /C118 0 /C119 0 /C118 0 /C117 0 /C117 0 /C117
MgO 7 /C167 7 /C185 8 /C129 7 /C187 8 /C117 7 /C188 5 /C176 5 /C162 5 /C168 5 /C153 5 /C134 5 /C158 5 /C136
CaO 10 /C162 10 /C139 10 /C140 10 /C133 10 /C142 10 /C141 10 /C127 10 /C136 10 /C114 9 /C196 8 /C153 9 /C142 9 /C105
Na2O1 /C194 1 /C192 1 /C189 1 /C194 1 /C189 1 /C196 2 /C113 2 /C120 2 /C108 2 /C108 3 /C107 2 /C194 3 /C101
K2O0 /C109 0 /C109 0 /C107 0 /C111 0 /C105 0 /C110 0 /C186 0 /C171 0 /C152 0 /C166 1 /C160 1 /C145 1 /C149
P2O5 0/C108 0 /C108 0 /C108 0 /C108 0 /C108 0 /C108 0 /C121 0 /C124 0 /C123 0 /C124 0 /C124 0 /C125 0 /C126
LOI 0 /C168 0 /C194 0 /C185 1 /C102 0 /C185 0 /C191 1 /C107 1 /C139 2 /C163 2 /C188 0 /C196 1 /C133 1 /C136
Total 99 /C188 99 /C179 100 /C109 100 /C113 99 /C178 100 /C105 99 /C134 99 /C195 99 /C190 99 /C134 100 /C113 100 /C111 100 /C101
Sc (ppm) 35 /C113 4 /C183 4 /C153 4 /C173 5 /C163 5 /C102 7 /C192 5 /C142 9 /C122 7 /C152 5 /C145 2 /C145 1 /C16
V 264 /C103 259 /C135 257 /C123 256 /C151 267 /C171 265 /C196 271 /C123 253 /C149 267 /C131 258 /C145 203 /C172 423 /C158 414 /C135
Rb 1 /C122 1 /C129 0 /C191 1 /C195 0 /C156 1 /C166 32 /C111 5 /C151 4 /C182 2 /C174 3 /C145 5 /C135 9 /C12
Sr 295 281 289 282 298 289 572 583 596 562 559 839 847Y1 4 /C111 3 /C161 3 /C141 3 /C191 3 /C181 3 /C142 0 /C162 1 /C172 3 /C162 1 /C193 0 /C123 1 /C103 0 /C19
Zr 28 /C102 7 /C152 7 /C112 7 /C162 8 /C102 7 /C195 5
/C186 1 /C106 7 /C186 5 /C11 125 115 124
Nb 1 /C123 1 /C121 1 /C121 1 /C123 1 /C128 1 /C121 2 /C191 2 /C173 3 /C102 3 /C126 4 /C186 4 /C181 4 /C115
Cs 0 /C1168 0 /C1176 0 /C1122 0 /C1269 0 /C1090 0 /C1215 0 /C1785 1 /C110 1 /C110 1 /C113 1 /C123 2 /C179 3 /C106
Ba 58 /C175 9 /C175 8 /C116 1 /C195 9 /C196 0 /C17 257 280 316 283 475 1012 1062
La 2 /C161 2 /C142 2 /C141 2 /C154 2 /C165 2 /C125 9 /C152 10 /C141 1 /C101 0 /C161 6 /C192 9 /C123 0 /C12
Ce 6 /C150 6 /C161 6 /C147 6 /C163 6 /C159 6 /C121 22 /C162 3 /C192 5 /C192 5 /C133 7 /C115 3 /C165 4 /C17
Pr 1 /C105 1 /C101 1 /C100 1 /C102 1 /C103 0 /C196 3 /C116 3 /C142 3 /C161 3 /C151 4 /C183 6 /C140 6 /C160
Nd 5 /C121 5 /C123 5 /C111 5 /C114 5 /C135 5 /C100 13 /C191 5 /C121 6 /C141 5 /C162 1 /C152 4 /C172 4 /C15
Sm 1 /C158 1 /C159 1 /C150 1 /C149 1 /C148 1 /C143 3 /C125 3 /C165 3 /C157 3 /C166 4 /C175 5 /C111 5 /C125
Eu 0 /C1652 0 /C1633 0 /C1628 0 /C1630 0 /C1607 0 /C1642 1 /C123 1 /C125 1 /C132 1 /C129 1 /C138 1 /C165 1 /C167
Gd 2 /C115 1 /C199 2 /C103 2 /C113 2 /C108 1 /C188 3 /C145 3 /C152 3 /C189 3 /C180 4 /C199 4 /C182 4 /C175
Tb 0 /C1352 0 /C1371 0 /C1358 0 /C1359 0 /C1371 0 /C1348 0 /C1542 0 /C1582 0 /C1601 0 /C1574 0 /C1807 0 /C1741 0 /C1710
Dy 2 /C115 2 /C114 2 /C116 2 /C111 2 /C110 2 /C110 3 /C127 3 /C136 3 /C156 3 /C128 4 /C170 4 /C149 4 /C135
Ho 0 /C1482 0 /C1483 0 /C1457 0 /C1501 0 /C1478 0 /C1458 0 /C1721 0 /C1714 0 /C1769 0 /C1742 1 /C104 0 /C1923 0 /C1924
Er 1 /C128 1 /C130 1 /C130 1 /C135 1 /C122 1 /C127 2 /C110 2 /C113 2 /C134 2 /C123 2 /C198 2 /C169 2 /C163
Tm 0 /C1216 0 /C1211 0 /C12106 0 /C1207 0 /C1202 0 /C1197 0 /C1298 0 /C1319 0 /C1334 0 /C1312 0 /C1412 0 /C1412 0 /C1434
Yb 1 /C132 1 /C129 1 /C119 1 /C131 1 /C135 1 /C136 1 /C198 2 /C103 2 /C136 2 /C128 2 /C173 2 /C180 2 /C186
Lu 0 /C1184 0 /C1192 0 /C1178 0 /C1189 0 /C1182 0 /C1193 0 /C1310 0 /C1332 0 /C1334 0 /C1327 0 /C1462 0 /C1469 0 /C1442
Hf 0 /C174 0 /C175 0 /C176 0 /C175 0 /C176 0 /C175 1 /C144 1 /C164 1 /C174 1 /C168 3 /C118 2 /C149 2 /C152
Ta 0 /C1066 0 /C1059 0 /C1062 0 /C1063 0 /C1062 0 /C1064 0 /C1149 0 /C1171 0 /C1158 0 /C1148 0 /C1227 0 /C1313 0 /C1269
Pb 2 /C155 2 /C138 2 /C170 2 /C155 2 /C156 2 /C144 2 /C116 2 /C194 3 /C169 3 /C139 4 /C197 8 /C193 9 /C104
Th 0 /C1239 0 /C1220 0 /C1217 0 /C1221 0 /C1214 0 /C1227 1 /C185 1 /C198 2 /C134 2 /C125 4 /C156 7 /C182 8 /C151
U0 /C1049 0 /C1055 0 /C1053 0 /C1056 0 /C1053 0 /C1055 0 /C1538 0 /C1553 0 /C1636 0 /C1567 1 /C129 1 /C190 2 /C108
(continued)2270 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Th/Ce and Th/Yb ratios than the Japan Trench sedi-
ments, as discussed below. These basalts often contain
fragments of granitic and sedimentary rock, signs thatthey could have been affected by crustal contamination.
In addition, some differentiated back-arc basalts
erupted at 10–8 Ma, 6–3 Ma and 2 /C15–0 Ma are character-
ized by high Th/Yb and La/Yb, consistent with crustal as-
similation. Lead isotopic analysis of the NEJ volcanicrocks in addition to Sr and Nd isotope study may clarify
the role of the crust in the petrogenesis of these rocks,
although Sr–Nd isotopes alone might not provide suchevidence ( Takahashi et al. , 2013 ). The above observa-
tions suggest that amongst the studied 35–0 Ma NEJ
basalts there are samples that may be affected by crus-
tal assimilation.Mantle composition beneath the NE Japan arc
independent of subduction modification
In general, one of the important geochemical charac-
teristics of island arc basalts is their depletion in HFSEcompared with LILE and LREE ( Fig. 10 ), a feature that is
not observed in MORB (e.g. Gill, 1981 ;Pearce et al. ,
2005 ) .E n r i c h m e n ti nL I L Ea n dL R E Ei ni s l a n da r cb a s -
alts may be caused by the addition of a subduction-
derived component to the arc magma source (e.g.Elliott et al. , 1997 ;Kimura & Yoshida, 2006 ;Hanyu
et al. , 2012 ), although heterogeneity of the mantle
wedge prior to the addition of a subduction-derivedcomponent may also result in diverse arc magma com-
positions (e.g. Linet al. , 1990 ;Woodhead et al. , 1993 ;
Hochstaedter et al. ,2 0 0 0 ,2001 ;Ishizuka et al. , 2003 ;Table 1 : Continued
Age: Quaternary (2 /C15–0 Ma)
Region: Back-arc
Basalt name: Sannome-gata
Sample no.: SMG01 SMG02 SMG04 SMG05
Age (Ma): 0 /C102 0 /C102 0 /C102 0 /C102
SiO 2(wt %) 49 /C111 48 /C174 50 /C109 50 /C122
TiO 2 0/C185 0 /C182 0 /C183 0 /C180
Al2O3 16/C131 15 /C164 16 /C118 16 /C142
FeO* 8 /C149 8 /C158 8 /C103 7 /C190
MnO 0 /C117 0 /C117 0 /C121 0 /C117
MgO 9 /C137 10 /C123 8 /C168 8 /C142
CaO 10 /C172 10 /C161 9 /C124 9 /C119
Na2O2 /C122 2 /C119 2 /C136 2 /C138
K2O1 /C100 0 /C195 1 /C129 1 /C108
P2O5 0/C116 0 /C116 0 /C117 0 /C118
LOI 1 /C180 1 /C147 2 /C118 2 /C142
Total 100 /C118 99 /C155 99 /C125 99 /C119
Sc (ppm) 39 /C113 8 /C163 3 /C173 2 /C12
V 270 268 248 239
Rb 21 /C172 1 /C103 2 /C132 5 /C19
Sr 496 490 473 490Y2 1 /C102 0 /C162 1 /C102 0 /C13
Zr 60 /C195 8 /C186 1 /C135 6 /C19
Nb 2 /C104 1 /C195 2 /C164 2 /C172
Cs 1 /C108 1 /C104 1 /C157 1 /C122
Ba 508 488 559 529La 14 /C181 3 /C191 5 /C151 5 /C19
Ce 28 /C182 6 /C143 1 /C162 9 /C18
Pr 3
/C190 3 /C165 4 /C100 4 /C110
Nd 16 /C171 6 /C111 6 /C171 7 /C18
Sm 3 /C167 3 /C167 3 /C174 3 /C178
Eu 1 /C116 1 /C105 1 /C107 1 /C109
Gd 3 /C176 3 /C183 3 /C178 3 /C178
Tb 0 /C1582 0 /C1551 0 /C1548 0 /C1584
Dy 3 /C149 3 /C145 3 /C129 3 /C142
Ho 0 /C1741 0 /C1724 0 /C1738 0 /C1707
Er 2 /C115 2 /C101 2 /C110 2 /C118
Tm 0 /C1299 0 /C1289 0 /C1296 0 /C1316
Yb 1 /C192 1 /C191 1 /C191 1 /C183
Lu 0 /C1314 0 /C1270 0 /C1304 0 /C1303
Hf 1 /C160 1 /C168 1 /C174 1 /C163
Ta 0 /C1116 0 /C1099 0 /C1133 0 /C1186
Pb 4 /C137 3 /C120 6 /C133 7 /C114
Th 4 /C128 3 /C192 5 /C125 5 /C133
U1 /C104 0 /C1966 1 /C124 1 /C125
*Total Fe as FeO.
LOI, loss of ignition; L., Lower.Journal of Petrology , 2015, Vol. 56, No. 11 2271Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Table 2 : Sr and Nd isotope data for selected basalts from the study area
Period Basalt name Sample no. Age (Ma)87Sr/86Sr87Rb/86Sr Sr I143Nd/144Nd147Sm/144Nd Nd I
Frontal arc
Early Miocene Takadate TA05 22 0 /C1704354 614 0 /C1046 0 /C1704339 614 0 /C1512843 614 0 /C11745 0 /C1512818 614
TA06 22 0 /C1704482 614 0 /C1079 0 /C1704457 614 0 /C1512753 614 0 /C11614 0 /C1512730 614
Tenmyosan TM01 20 0 /C1704301 614 0 /C1062 0 /C1704013 614 0 /C1512877 614 0 /C11800 0 /C1512853 614
TM13 20 0 /C1704234 614 0 /C1037 0 /C1704224 614 0 /C1512748 614 0 /C11926 0 /C1512723 614
TM14 20 0 /C1704235 614 0 /C1053 0 /C1704220 614 0 /C1512883 614 0 /C11840 0 /C1512859 614
Middle Miocene Rokugo HG15 15 0 /C1704278 614 0 /C1031 0 /C1704271 614 0 /C1512854 614 0 /C11761 0 /C1512837 614
Nakuidake 0604 15 0 /C1704120 614 0 /C1067 0 /C1704105 614 0 /C1512706 614 0 /C11673 0 /C1512690 614
Washinosu WA06 12 0 /C1704816 614 0 /C1037 0 /C1704810 614 0 /C1512740 613 0 /C11591 0 /C1512728 613
0628E 12 0 /C1705065 614 0 /C1105 0 /C1705047 614 0 /C1512673 612 0 /C11549 0 /C1512661 612
Late Miocene Joge 71 8 /C120 /C1704363 614 0 /C1029 0 /C1704359 614 0 /C1512892 614 0 /C11691 0 /C1512883 614
SI14 8 /C120 /C1704382 614 0 /C1020 0 /C1704379 614 0 /C1512836 614 0 /C12173 0 /C1512824 614
J18 8 /C120 /C1704414 614 0 /C1041 0 /C1704409 614 0 /C1512776 614 0 /C11960 0 /C1512765 614
Mitaki TA1 7 /C190 /C1704294 614 0 /C1059 0 /C1704287 614 0 /C1512859 613 0 /C12076 0 /C1512849 613
TA2 7 /C190 /C1704299 613 0 /C1061 0 /C1704293 613 0 /C1512860 618 0 /C12139 0 /C1512849 618
TA3 7 /C190 /C1704298 614 0 /C1049 0 /C1704293 614 0 /C1512887 614 0 /C12043 0 /C1512876 614
Pliocene Kurohanayama 090901 5 /C160 /C1704213 614 0 /C1030 0 /C1704210 614 0 /C1512725 614 0 /C11881 0 /C1512718 614
090902 5 /C160 /C1704268 614 0 /C1046 0 /C1704264 614 0 /C1512821 614 0 /C11883 0 /C1512814 614
Araya AR01 3 0 /C1704117 614 0 /C1012 0 /C1704117 614 0 /C1512844 613 0 /C12168 0 /C1512839 613
AR02 3 0 /C1704125 614 0 /C1022 0 /C1704124 614 0 /C1512888 614 0 /C12136 0 /C1512883 614
Hakaseyama 90403 2 /C180 /C1703944 613 0 /C1049 0 /C1703942 613 0 /C1512834 614 0 /C11635 0 /C1512831 614
420 2 /C180 /C1703925 612 0 /C1112 0 /C1703921 612 0 /C1512869 614 0 /C11514 0 /C1512866 614
Inaniwadake IN10 2 /C170 /C1704157 614 0 /C1010 0 /C1704157 614 0 /C1512668 614 0 /C12177 0 /C1512664 614
IN13 2 /C170 /C1704355 614 0 /C1022 0 /C1704354 614 0 /C1512855 613 0 /C11961 0 /C1512852 613
Quaternary Funagata FU02 0 /C190 /C1704167 614 0 /C1039 0 /C1704166 614 0 /C1512849 614 0 /C11900 0 /C1512848 614
FU05 0 /C190 /C1704109 614 0 /C1022 0 /C1704109 614 0 /C1512803 614 0 /C11794 0 /C1512802 614
FU06 0 /C190 /C1704119 614 0 /C1022 0 /C1704118 614 0 /C1512824 614 0 /C11717 0 /C1512823 614
Iwate IW03 0 /C110 /C1704190 614 0 /C1009 0 /C1704190 614 0 /C1512786 614 0 /C11775 0 /C1512786 614
IW05 0 /C110 /C1704179 614 0 /C1005 0 /C1704179 614 0 /C1512769 614 0 /C11673 0 /C1512769 614
IW06 0 /C110 /C1704171 613 0 /C1017 0 /C1704171 613 0 /C1512814 614 0 /C11730 0 /C1512814 614
Back-arc
Early Miocene L. Fukuyama BMT03* 24 /C120 /C1704508 614 0 /C1351* 0 /C1704388 614 0 /C1512688 614 0 /C11428 0 /C1512665 614
YO10* 24 /C120 /C1704644 614 0 /C1095* 0 /C1704612 614 0 /C1512622 614 0 /C11316 0 /C1512601 614
Late Miocene Aosawa NM03 10 /C140 /C1703778 613 0 /C1070 0 /C1703767 613 0 /C1512893 610 0 /C11580 0 /C1512882 610
NM24 10 /C140 /C1703496 612 0 /C1023 0 /C1703492 612 0 /C1512963 610 0 /C11620 0 /C1512952 610
NM42 10 /C140 /C1703506 614 0 /C1089 0 /C1703492 614 0 /C1512939 607 0 /C11655 0 /C1512928 607
Aizome AZ01 10 /C120 /C1703755 614 0 /C1167 0 /C1703731 614 0 /C1512907 614 0 /C11530 0 /C1512897 614
AZ02 10 /C120 /C1703967 614 0 /C1126 0 /C1703949 614 0 /C1512866 613 0 /C11576 0 /C1512855 613
AZ03 10 /C120 /C1703960 614 0 /C1101 0 /C1703946 614 0 /C1512866 614 0 /C11557 0 /C1512855 614
AZ04 10 /C120 /C1703534 614 0 /C1088 0 /C1703521 614 0 /C1512921 613 0 /C11377 0 /C1512912 613
Pliocene Manaitayama MIY06 3 /C160 /C1703738 614 0 /C1050 0 /C1703735 614 0 /C1512861 612 0 /C11449 0 /C1512858 612
Quaternary Asakusa 54109 1 0 /C1703946 614 0 /C1072 0 /C1703945 614 0 /C1512788 613 0 /C11319 0 /C1512787 613
61106 1 0 /C1703907 614 0 /C1117 0 /C1703905 614 0 /C1512809 610 0 /C11421 0 /C1512808 614
Chokai CH02 0 0 /C1702970 614 0 /C1224 0 /C1702970 614 0 /C1512860 614 0 /C11334 0 /C1512860 614
Kampu 81710 0 /C110 /C1703209 614 0 /C1191 0 /C1703208 614 0 /C1512934 614 0 /C11250 0 /C1512934 614
82008 0 /C110 /C1703265 611 0 /C1202 0 /C1703264 611 0 /C1512930 609 0 /C11295 0 /C1512930 609
Sannome-gata SMG01 0 /C102 0 /C1703197 614 0 /C1127 0 /C1703197 614 0 /C1512912 613 0 /C11329 0 /C1512912 613
SMG02 0 /C102 0 /C1703180 612 0 /C1124 0 /C1703180 612 0 /C1512953 614 0 /C11375 0 /C1512953 614
SMG04 0 /C102 0 /C1703227 614 0 /C1197 0 /C1703227 614 0 /C1512956 614 0 /C11353 0 /C1512956 614
SrI, initial87Sr/86Sr ratio; Nd I, initial143Nd/144Nd ratio. L., Lower.
*Contents of Rb, Sr, Nd and Sm are reported in SD Electronic Appendix Table A2.2272 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Fig. 3. Variation of selected trace element concentrations and trace element ratios vs MgO (wt %) for early Oligocene (35–32 Ma)
and early Miocene (24–18 Ma) basalts (MgO >5 wt %).Journal of Petrology , 2015, Vol. 56, No. 11 2273Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Caulfield et al. , 2008 ). HFSE such as Nb, Ta, Zr and Hf
in subduction-related rocks are widely regarded as
fluid- and melt-immobile (e.g. Mu¨ nker et al. , 2004 ;
Pearce et al. , 2005 ;Zack & Timm, 2007 ), and thus theabundances and ratios of these elements are suitable
to explore the mantle source composition of island arc
magmas, and to what extent the source has been
melted.Fig. 4. Variation of selected trace element concentrations and trace element ratios vs MgO (wt %) for middle Miocene (16–12 Ma)
basalts (MgO >5 wt %).2274 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Importantly, experimentally determined distribution
coefficients for common upper mantle assemblages
such as spinel lherzolite show that HFSE incompatibility
is Nb>Ta>Zr>Hf (Mu¨ nker et al. , 2004 , and referencestherein). Thus, melt extraction should result in a mantle
residue having Nb/Ta ratios less than chondrite (17 /C16;
Weyer et al. , 2002 ) at low Nb abundances and low Zr/Hf
ratios, and this feature would be observed in arcFig. 5. Variation of selected trace element concentrations and trace element ratios vs MgO (wt %) for late Miocene (10–8 Ma) basalts
(MgO>5 wt %).Journal of Petrology , 2015, Vol. 56, No. 11 2275Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
magma sources that are more depleted than the sur-
rounding upper mantle ( Elliott, 2003 ).
Plots of Zr versus Zr/Hf ( Fig. 12 a and b) include the
MORB field. In general, the NEJ basalts produced at 35–32 Ma, 24–18 Ma, 16–12 Ma, 10–8 Ma, 6–3 Ma and 2 /C15–
0 Ma show an increase in Zr/Hf with increasing Zr from
the frontal arc to the back-arc ( Fig. 12 a), and although
some analyses of back-arc basalts plot outside theFig. 6. Variation of selected trace element concentrations and trace element ratios vs MgO (wt %) for Pliocene (6–3 Ma) basalts
(MgO>5 wt %).2276 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
MORB field, most frontal- and back-arc basalt data over-
lap with this field, as do most of the 16–12 Ma basalts
(Fig. 12 b). These features indicate that the composition
of the mantle source of the NEJ basalt magmas prior tothe addition of a subduction-derived component
approximated DMM.
As discussed below, the depth of segregation from
the mantle for most Oligocene to Quaternary primaryFig. 7. Variation of selected trace element concentrations and trace element ratios versus MgO (wt %) for Quaternary (2 /C15–0 Ma)
basalts (MgO >5 wt %).Journal of Petrology , 2015, Vol. 56, No. 11 2277Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
magmas to the NEJ basalts has been estimated to be
less than about 75 km ( Tatsumi et al. , 1983 ,1994 ;
Kimura et al. , 2009 ;Sato et al. , 2013 ,2014 ;Shuto et al. ,
2013 ;Kuritani et al. , 2014 b, and references therein),corresponding to the stability field of spinel ( Robinson
& Wood, 1998 ). This is consistent with a previously pro-
posed petrological model for an upper mantle com-
posed of spinel lherzolite beneath the IchinomegataFig. 8. Age vs selected trace element concentrations and ratios for 35–0 Ma NEJ basalts. For samples with more than one analysis,
average values have been used. Values of N-MORB, E-MORB and OIB are from Sun & McDonough (1989) , and DMM values are
from Workman & Hart (2005) .2278 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
volcano in the NE Japan back-arc region, which is
based on the petrology of ultramafic to mafic xenolithsexhumed by the volcanic rocks ( Takahashi, 1986 ). Thus,
spinel peridotite may be widely distributed in the man-
tle wedge beneath the NE Japan arc. We use the modal
assemblage for a spinel peridotite (olivine, orthopyrox-ene, clinopyroxene and spinel in the proportions0/C157:0/C128:0/C113:0/C102) to model the DMM mantle ( Fig. 13 ).
Different degrees of melting of DMM: a possible
mechanism for the generation of NEJ basalt
magmas
Plots of Nb vs Nb/Yb ( Fig. 13 a and b) include a melting
curve for DMM ( Workman & Hart, 2005 ) calculated
using the batch partial melting equation of Shaw (1970) .
In these diagrams are shown the average Nb contents
and Nb/Yb ratios of 43 basalts ( Figs 3 –7), except for the
16–12 Ma frontal-arc Ishimoriyama and Washinosu bas-alts, for which the values for the least differentiatedsamples are used. Both the Nb content and Nb/Yb ratioin basaltic melts should increase with olivine fraction-ation, but there is little change (especially for Nb/Y) for
most basalts over a given MgO range except for some
basalts such as the Ishimoriyama basalt, earlyOligocene Kamo basalt, early Miocene Lower
Fukuyama basalt, late Miocene Aizome basalt andQuaternary Sannome-gata basalt ( Figs 3 –7); thus a
comparison of Nb and Nb/Yb for the respective basaltic
magmas can be made using average element contents
and ratios.
Most NEJ basalt data plot on or near a modal batch
partial melting line ( Fig. 13 a). In more detail, frontal-arc
basalts result from between about 5 and 18% melting(24–18 Ma, 5–10%; 10–8 Ma, 15%; 6–3 Ma, 8–18%,
2/C15–0 Ma, 9–12%), whereas back-arc basalts have lower
values between 1 /C15 and 6%. Although such a distinction
is also seen for frontal-arc and back-arc basalts at16–12 Ma ( Fig. 13 b), the degree of melting for the fron-
tal-arc Ishimoriyama magma is within the melting rangeof the back-arc magmas, and that for the back-arcTappizaki magma is similar to that of most frontal-arc
magmas. The higher degrees of partial melting of a
DMM source for frontal-arc basalts are consistent withthe conclusions of Kimura et al. (2009) and Kuritani
et al. (2014 b), who estimated higher degrees of melting
for Quaternary frontal-arc basalt magmas comparedwith back-arc magmas based on mass-balance model-
ing and thermodynamic analysis of the petrological
characteristics of the basalts. The data for theFig. 9. Age vs Nb/Yb (a) and Zr/Y (b) for middle Miocene (16–12 Ma) basalts. All the analyzed data are plotted.Journal of Petrology , 2015, Vol. 56, No. 11 2279Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Inaniwadake and Araya basalts indicate a source that is
slightly depleted in Nb compared with DMM. The vari-ation of Nb/Y with age ( Fig. 8 c) supports Nb depletion inthe two basalt magma sources. Additionally, it is likely
that the source for these two basalt magmas wasdepleted in Zr relative to DMM, consistent with lowerFig. 10. N-MORB-normalized trace element patterns for the NEJ basalts. (a) Early Oligocene (35–32 Ma), (b) early Miocene (24–
18 Ma), (c) middle Miocene (16–12 Ma), (d) late Miocene (10–8 Ma), (e) Pliocene (6–3 Ma), and (f) Quaternary (2 /C15–0 Ma). N-MORB
values are from Sun & McDonough (1989) . The trace element data plotted within gray bands show distinct differences in the
degrees of depletion for Nb, Ta, Zr and Hf between the frontal-arc and back-arc basalts erupted at 24-18 Ma, 10-8 Ma, 6-3 Ma and2.5-0 Ma, but no clear differences between the frontal-arc and back-arc basalts at 16-12 Ma.2280 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Zr/Y and Zr/Lu ratios in both basalts ( Figs. 8 d and e).
Based on whole-rock and mineral chemistry, Yasui &
Yamamoto (2006) showed that the Inaniwadake magma
had a low oxygen fugacity and H 2O content and waserupted at a temperature of between 1053 and 1203/C14C. In
contrast to these two basalts, some basalts (early
Oligocene Matsue basalt, an d early Miocene Aonaegawa,
Miyano and Nomuragawa basalts) plot below the partialFig. 11. C1 chondrite-normalized REE patterns for the NEJ basalts. (a) Early Oligocene (35–32 Ma), (b) early Miocene (24–18 Ma), (c)
middle Miocene (16–12 Ma), (d) late Miocene (10–8 Ma), (e) Pliocene (6–3 Ma), and (f) Quaternary (2 /C15–0 Ma). C1 chondrite values
are from Sun & McDonough (1989) .Journal of Petrology , 2015, Vol. 56, No. 11 2281Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
melting line ( Fig. 13 a), indicating a source that is slightly
enriched in Nb compared with DMM.
Relationships between degrees of partial melting
of DMM and segregation depth for the NEJ
basalt magmas
Tatsumi et al. (1994) used multi-saturation experiments
to show similar segregation pressure of 1 /C12a n d1 /C14G P a
for the primary magmas of the middle Miocene Tomariand Ryozen basalts from the frontal-arc region of the NEJapan arc. Shuto et al. (2013) arrived at similar results of
1 GPa for the Tomari magma and 1 /C15 GPa for the Ryozen
magma using the olivine maximum fractionation model,and argued that the middle Miocene (16–12 Ma) frontal-arc primary basalt magmas may have segregated froman/C2420 km thick peridotite region (between 1 /C10a n d
1/C15 GPa) of the uppermost mantle, lying just below the
Moho discontinuity beneath the NEJ frontal-arc region,
which extended over a distance of about 500 km ( Fig. 1 ).
The experiments of Tatsumi et al. (1994) showed that
22 Ma Mg-rich basalt magmas from the back-arc regionnear Akita segregated from the mantle at depths of30–40 km. They argued that there was no across-arcvariation in the depth of magma segregation in the
Miocene NE Japan arc, and that the depth of magma
segregation has remained constant since 20 Ma in thefrontal-arc region, whereas it changed from 30–40 km to>60 km beneath the back-arc region.
Based on the marked similarities in major and trace
element chemistry and normative mineralogy between
two early Oligocene and early Miocene basalts from
southern Okushiri Island (the 34 /C14 Ma Matsue basaltand the 22 Ma Aonaegawa basalt; Fig. 1 ) and the
Quaternary Daisen basalt in the SW Japan arc, Satoet al. (2013, 2014 ) suggested a segregation depth for all
three primary magmas of about 60 km ( Tamura et al. ,
2000 ). Therefore, it is more prudent to restrict the obser-
vation of Tatsumi et al. (1994) about the lack of across-
arc variation in magma segregation depth to the middleMiocene, rather than the entire Miocene.
I tc a nb ea s s u m e dt h a th i g hd e g r e e so fm e l t i n ga ts h a l -
low depths of 30–50 km in the mantle wedge beneath thefrontal arc persisted from 22 Ma to the present. In theback-arc, lower degrees of partial melting from 35 Ma to
the present took place at deeper levels (60–75 km) except
for the middle Miocene ba ck-arc Tappizaki magmas,
which resulted from similar degrees of melting to the fron-tal-arc magmas during the same period.
Identification of the subduction components
For a DMM mantle source that is melted in the spinelperidotite stability field, the lower Nb/Yb of the frontal-arc basalts at 24–18 Ma, 10–8 Ma, 6–3 Ma and 2 /C15–0 Ma,
and most of the 16–12 Ma frontal-arc basalts, comparedwith the back-arc basalts of the same ages ( Figs 3 c, 5c,
6c, 7 c and 8 c) suggests higher depletion and/or higherdegrees of melting. However, assuming pure DMM as
the mantle source composition, this is inconsistent with
Nb/Yb versus Ba/Yb, Th/Yb and La/Yb variations of 35–0 Ma NEJ basalts, which plot in the field above theMORB–OIB mantle array ( Fig. 14 a–c).
These features and the observations from the plot of
Nb/Yb versus Nb ( Fig. 13 ) indicate that the composition
of the NEJ basalt magmas may be affected by the add-
ition of subduction components. Barium, but not ThFig. 12. Zr vs Zr/Hf for the NEJ basalts. (a) Early Oligocene (35–32 Ma), early Miocene (24–18 Ma), late Miocene (10–8 Ma), Pliocene
(6–3 Ma), and Quaternary (2 /C15–0 Ma), and (b) middle Miocene (16–12 Ma). Field of MORB is compiled from the petrological database
PetDb (Conny & Kerstin, 2012).2282 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
and LREE, is readily partitioned into aqueous fluids
released from the subducting slab at relatively low tem-perature and pressure, whereas Ba, Th and LREE aremore likely to be extracted from the subducting slab atrelatively high temperature and pressure as a sediment-
derived melt ( Pearce et al. , 2005 , and references
therein). Thus, Th/Nb, Ba/Nb and Ba/Th can be used toidentify the nature of the subduction component.Pearce et al. (2005) proposed that Ba/Th and Th/Nb are
proxies for the addition of a shallow subduction compo-nent (hydrous fluid) and deep subduction component(sediment melt), respectively, whereas Ba/Nb is a proxy
for addition of a total subduction component. Th/Ce is
also sensitive to deep subduction addition becausesubducting sediment generally has a higher Th contentthan the mantle, so that the Th/Ce ratio is elevated com-pared with the mantle ( Plank & Langmuir, 1998 ).
Figure 14 b and c indicates that the main subduction
component added to the DMM source is sediment melt
(i.e. a deep subduction component). However, Fig. 14 asuggests addition of a total subduction component to
DMM, and thus it is necessary to identify whether enrich-ment in Ba/Yb for these basalts was caused by additionof a hydrous fluid (shallow subduction component) and/or a sediment melt. Back-arc lavas from the Izu–Bonin
arc ( Ishizuka et al. , 2003 ) and two types of primitive bas-
alt (COB1 and COB2) from the Pagan volcano within theMariana arc volcanic front ( Tamura et al. ,2 0 1 4 ) are simi-
lar to the NEJ basalts ( Fig. 14 a–c).
In a Th/Zr vs Ba/Zr plot ( Fig. 15 a), most samples of
the 35–0 Ma NEJ basalts form a broad trend in whichthe Th/Zr ratio increases with increasing Ba/Zr. This
trend is similar to a mixing line between Pacific MORB
and bulk sediment from the Japan Trench ( Plank &
Langmuir, 1998 ; hereafter referred to as Japan Trench
Sediment; JTS). In terms of Th/Ce versus Th/Yb(Fig. 15 b), most NEJ basalts plot subparallel to the
Pacific MORB–JTS mixing line. Thus, it is likely that the
main subduction component added to the mantle
source of the 35–0 Ma NEJ basalt magmas is sedimentFig. 13. Nb/Yb vs Nb for the NEJ basalts. (a) Early Oligocene (35–32 Ma), early Miocene (24–18 Ma), late Miocene (10–8 Ma),
Pliocene (6–3 Ma), and Quaternary (2 /C15–0 Ma), and (b) middle Miocene (16–12 Ma). In these diagrams, average concentrations and
ratios of these elements in each of the 45 basalts shown in Figs 3 –8are plotted. Continuous line A in (a) and (b) shows batch partial
melting of DMM with the modal assemblage of a spinel peridotite composed of olivine, orthopyroxene, clinopyroxene and spinel(0/C157:0/C128:0/C113:0/C102;Workman & Hart, 2005 ). Mineral–melt partition coefficients are 0 /C100017 (Nb) and 0 /C1056 (Yb) for olivine, 0 /C10028
(Nb) and 0 /C10891 (Yb) for orthopyroxene, 0 /C1008 (Nb) and 0 /C15819 (Yb) for clinopyroxene, and 0 /C108 (Nb) and 0 /C10076 (Yb) for spinel.
Olivine partition coefficients are from Zanetti et al. (2004) , orthopyroxene and clinopyroxene partition coefficients are from McDade
et al. (2003) , and spinel partition coefficients are from Horn et al. (1994) . Data for N-MORB and E-MORB are from Sun &
McDonough (1989) , and those for Pacific MORB are from Arevalo & McDonough (2010) .Journal of Petrology , 2015, Vol. 56, No. 11 2283Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Fig. 14. (a) Nb/Yb vs Ba/Yb, (b) Nb/Yb vs Th/Yb and (c) Nb/Yb vs La/Yb. Data sources: N-MORB, E-MORB and OIB are from Sun &
McDonough (1989) ; JTS (bulk compositions of sediment column subducting at the Japan Trench) from Plank & Langmuir (1998) ;
fields of COB1 and COB2 (two types of primitive basalt, clinopyroxene–olivine basalt 1 and clinopyroxene–olivine basalt 2, from thePagan volcano at the Mariana arc volcanic front) from Tamura et al. (2014) ; field of Izu–Bonin back-arc lavas from Ishizuka et al.
(2003) . MORB-OIB array in (b) is from Pearce (2008) . DMM is from Workman & Hart (2005) .2284 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
melt. Similar arguments have been made for subduc-
tion components added to the source mantle beneaththe Izu–Bonin back-arc region and the Pagan volcano.The Izu–Bonin back-arc lavas have high Th/Zr and lowBa/Zr, and form a trend subparallel to a MORB–bulk
sediment mixing line ( Ishizuka et al. , 2003 ,fig. 5 ). They
also have significantly higher Th/Ce compared withMORB ( Fig. 15 b). This led Ishizuka et al. (2003) to con-
clude that subducted bulk sediment is an importantcomponent in the genesis of these lavas. A similar ex-planation has been applied to some primary basalts
from Pagan, which have higher Th/Nb and LREE,
whereas others are depleted in LREE and have lowerTh/Nb, indicative of a hydrous fluid ( Tamura et al. ,
2014 ). Therefore, we conclude from Fig. 15 a and b that
enrichment in Ba/Yb for the NEJ basalt magmas(Fig. 14 a) cannot be explained by addition of slab-
derived fluid to the source mantle, but can be explained
by sediment melt addition.
The unusually higher Ba/Zr, Th/Zr, Th/Ce and Th/Yb of
16 samples compared with many other NEJ basalts(enclosed with a dashed line in Fig. 15 a and b) may re-
flect not only subduction inputs to the source mantle butalso assimilation of crustal sediment. This is consistent
with the diversity of lithic fragments (including volcanic
and granitic rocks and siltstones) found in theQuaternary Sannome-gata basalt, which has affected the
major and trace element chemistry of these rocks(Yoshinaga & Nakagawa, 1999 ;Kuritani et al. , 2014 b). It
is difficult quantitatively to distinguish the contributionsof crustal sediment-derived trace elements from the
total amount of sediment-derived trace elements in the
basalts. Here, we have evaluated these 16 samples[four samples for the Tomari basalt (middle Miocene),one sample for Kaminokuni basalt (late Miocene),three samples for Toppu basalt (Pliocene), one samplefor Kamuiyama basalt (Pliocene), one sample
for Hakaseyama basalt (Pliocene), four samples for
Sannome-gata basalt (Quaternary) and two samples forKampu basalta (Quaternary)], most of which have higherBa/Zr and Th/Ce than JST ( Fig. 15 a and b), as those that
have undergone assimilation of crustal sediments.
The nature of across-arc variations in Sr and Nd
isotope composition for the 35–0 Ma NEJ basalts;
implications for a two-layered structure of themantle wedge
A compilation of Sr and Nd isotope data for 35–0 Ma
NEJ basalts ( Shuto et al. , 2013 ) has convincingly shown
their across-arc variations. Most of the basalts erupted
along the NE Japan arc between 35 and 15 Ma haveFig. 15. (a) Th/Zr vs Ba/Zr and (b) Th/Yb vs Th/Ce for the NEJ basalts. Several NEJ basalt samples within the field enclosed by a
long-dashed line in (a) and (b) show the effects of crustal assimilation (see text). JTS in (a) and (b) are the same as in Fig. 14 . Pacific
MORB in (a) and (b) is from Arevalo & McDonough (2010) . The fields of COB1, COB2 and Izu–Bonin back-arc lavas in (a) and (b) are
the same as in Fig. 14 . Continuous line in (a) and (b) shows bulk mixing between Pacific MORB and JTS.Journal of Petrology , 2015, Vol. 56, No. 11 2285Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
isotopically enriched compositions (Sr Ibetween 0 /C17040
and 0 /C17065 and Nd Ibetween 0 /C15124 and 0 /C15128), except
for basalts on Okushiri Island (Sr I0/C17035–0 /C17040), which
were erupted at 34 /C14 Ma. Most back-arc basalts younger
than 15 Ma have more depleted isotopic compositions
(SrI¼0/C17029–0 /C17040 and Nd I¼0/C151284–0 /C151312). In
marked contrast, the frontal-arc basalts remain enrichedfrom 22 Ma to the present. Our new data are consistentwith these observations.
We reconstructed Sr
Iversus Nd Idiagrams for 35–
0 Ma back-arc basalts ( Fig. 16 a) and 24–0 Ma frontal-arc
basalts ( Fig. 16 b) using the data of Shuto et al. (2013)
and this study. Most back-arc basalts younger than
16 Ma plot in a slightly enriched region compared withPacific Ocean MORB in terms of Sr
I–Nd I(Fig. 16 a). In
contrast, back-arc basalts older than 18 Ma have Sr Iand
NdIvalues that are even more enriched than the
younger basalts. As a result, Pacific Ocean MORB and
younger and older back-arc basalts form an array ex-
tending toward JTS ( Fig. 16 a). In contrast, frontal-arc
basalts older than 18 Ma overlap with basalts youngerthan 16 Ma. These frontal-arc basalts partly overlapwith younger back-arc basalts, but most lie near orwithin the region of older back-arc basalts. This indi-
cates a similar, enriched Sr and Nd isotope signature in
the back-arc basalts older than 18 Ma and most of the24–0 Ma frontal-arc basalts. Hanyu et al. (2006) have re-
ported Sr, Nd and Hf isotopic compositions for 25 Maand younger frontal-arc basalts similar to those studiedhere, and demonstrated that these basalts show littlechange in Sr and Nd isotopic compositions according
to age (similar to what is reported here), but exhibit
clear temporal variations in Hf isotopic ratios from theearly Miocene to the present. They have attributed thisto a change in the composition of the metasomatizingagent derived from the subducted slab. Furthermore, al-though Sr and Nd isotopes do not support crustal as-
similation for the Quaternary frontal basalts such as the
Azuma, Iwate, Funagata, and Akita-koma volcanoes,samples from Azuma volcano have more radiogenic Pbisotopic compositions than other basalts ( Shibata &
Nakamura, 1997 ;Takahashi et al. , 2013 ), indicative of
crustal assimilation. This shows that a combined iso-topic approach (i.e. Sr–Nd–Pb–Hf) is required to confirm
the influence of crustal assimilation versus slab input.
The temporal changes in Sr and Nd isotope compos-
itions for the NEJ basalts have led some workers to pro-pose that the mantle wedge beneath the NE Japan arcwas composed of two parts, an isotopically enrichedsubcontinenal lithospheric mantle and a depleted as-
thenospheric mantle ( Nohda et al. , 1988 ;Ohki et al. ,
1993 ). Additional insight into this two-layered structure
of the mantle wedge can be obtained by the temporalchange in isotope chemistry of basaltic rocks recoveredfrom the Yamato Basin in the Japan Sea. Cousens et al.
(1994) reported Sr and Nd isotope data for basalts col-
lected from Ocean Drilling Program (ODP) Sites 794,
795, and 797. These basalts were classified into younger
depleted and older less depleted groups (D and LDgroups) ( Shuto et al. , 2004 ;Nohda, 2009 ). The D group
younger than about 15 Ma is more depleted in terms ofSr and Nd isotope ratios, whereas the LD group olderthan 15 Ma is less depleted. It is worth noting that the Srand Nd isotope ratios of the D group are similar to those
of the NEJ back-arc basalts younger than 16 Ma, and
those of the LD group lie within the Sr and Nd isotoperange of the NEJ back-arc basalts older than 18 Ma.Thus, this temporal isotopic change in the back-arcbasin basalts is roughly consistent with that shown bythe 35–0 Ma NEJ back-arc basalts, indicating that themantle source for basalt magmas beneath the Japan
Sea and the back-arc side of the NE Japan arc was
replaced by isotopically depleted asthenospheric man-tle, which upwelled at around 15–16 Ma ( Shuto et al. ,
2004 ;Nohda, 2009 ).
It has been discussed that the change in the isotopic
compositions of the NEJ back-arc basalts is closely
linked with the upwelling of hot depleted asthenospheric
mantle during the opening of the Japan Sea ( Nohda
et al. ,1 9 8 8 ;Tatsumi et al. ,1 9 8 9 ;Shuto et al. ,2 0 0 4 ), which
may also have implications for the role of hot depletedasthenospheric upwelling in the production of the 35–0 Ma NEJ basalt magmas. In this context, Nohda et al.
(1988) argued that the relative contribution of the two
mantle components controlled the Sr and Nd isotopiccompositions of the basalt magmas; asthenosphere-derived magmas would undergo significant interactionwith the lithospheric mantle prior to back-arc opening,but less interaction after 15 Ma owing to thinning of thelithospheric mantle as a result of back-arc opening. This
is a plausible process to account for the temporal change
in the isotopic composition of the 35–0 Ma NEJ back-arcbasalts. Accordingly, the most enriched isotopic valuesfor the subcontinental lithospheric mantle are estimatedto be about Sr
I¼0/C17055–0 /C17065 and Nd I¼0/C15125–0 /C15126
at 15 Ma ( Fig. 16 a and b).
An important question is how asthenospheric upwell-
ing affected the genesis of the NEJ frontal-arc magmas.With regard to this, Shuto et al. (2013) suggested that
frontal-arc basalts were produced by a similar process tothat responsible for back-arc basalts described above,and that the upwelling of depleted asthenosphere dis-cussed here could have reached the upper mantle be-
neath the frontal-arc region by 22–16 Ma (pre- to syn-
opening stage of the Japan Sea), resulting in a highergeothermal gradient in the mantle wedge ( Shuto et al. ,
2013 ,fig. 20 ). In such a situation, asthenospheric melting
would have occurred and subsequent interaction be-tween the asthenosphere-derived melts and the subcon-
tinental lithospheric mantle would result in more
enriched frontal-arc basalts from 20 Ma to the present.
Variability in source mantle compositions for the
NEJ basalts
We have argued that variations in abundances and
ratios of incompatible elements for most of the NEJ
basalts, except for several older (older than 18 Ma)2286 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
back-arc basalts and some 6–3 Ma frontal-arc basalts,
can be accounted for by variable degrees of melting of
a DMM source modified by addition of subducted
sediment melts (referred to as M-DMM). However, the35–32 Ma and 24–18 Ma back-arc basalts, such as the
Matsue, Aonaegawa, Miyano, and Nomuragawa bas-
alts, cannot be produced from such a mantle source,
but may be generated from enriched subcontinentalFig. 16. SrIvs Nd Idiagram for (a) NEJ back-arc basalts and (b) NEJ frontal-arc basalts. Sr and Nd isotopic data for these basalts are
from Shuto et al. (2013) and this study. MORB data from the Pacific Ocean are from Conny & Kerstin (2012) . Calculated mixing line
between DMM and JTS is based on the equation of Langmuir et al. (1978) . DMM values from Workman & Hart (2005) ; Sr isotopic
ratio 0 /C170263, Nd isotopic ratio 0 /C151313, Sr 7 /C1664 ppm, Nd 0 /C1581 ppm. JTS values (at 15 Ma) from Plank & Langmuir (1998) ;S r I
0/C171079, Nd I0/C151233, Sr 87 ppm, Nd 22 /C103 ppm. Numbers on the mixing line indicate the per cent contribution from JTS. HIMU, EM
I, EM II and primitive mantle values are from Hart et al. (1986) .Journal of Petrology , 2015, Vol. 56, No. 11 2287Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
lithosphere. Similarly, some frontal-arc 6–3 Ma basalts
such as the Inaniwadake and Araya basalts wereformed from a source mantle that is slightly depleted interms of HFSE compared with DMM ( Fig. 13 )
The M-DMM source may mirror the composition of
depleted asthenosphere from which isotopically more
depleted basalts (most of the back-arc basalts youngerthan 16 Ma) were derived. In contrast, melting of as-thenospheric mantle and subsequent interaction be-tween the resultant magmas and the subcontinenallithospheric mantle would result in isotopically moreenriched basalts, including pre-18 Ma back-arc basalts
and post-22 Ma frontal-arc basalts. Although there are
large differences in the Sr and Nd isotope compositionsof more depleted and more enriched NEJ basalts, mostof the more depleted basalts and some of the more en-riched basalts plot on or near an identical batch partialmelting curve for DMM in terms of Nb/Yb versus Nb
(Fig. 13 ). This leads us to assume that more depleted as-
thenosphere-derived basalt magma interacted with en-riched lithospheric mantle, resulting in modifiedmagmas characterized by enriched Sr and Nd isotopesignatures. However, both the concentrations and ratiosof HFSE and HREE were unaffected because of their
melt immobility.
Estimating the amount of the subduction
component added to the mantle wedge beneath
the NE Japan arc
It is possible that the Sr and Nd isotope composition of
the depleted asthenosphere could have resulted frommixing of DMM and subducting sediments, and that ofthe enriched lithosphere also could have been affectedby subducting sediments. Basaltic volcanism in the NEJapan arc has continued inter mittently for a relatively
long period (35 Ma to the present), with the possibility
that sediment melt-mobile elements have been
released from the subducting slab into the mantlewedge over this time period. Although the subductedsediments are diverse in terms of lithology, Sr and Ndisotopic ratios, and Sr and Nd abundances ( Plank &
Langmuir, 1998 ), we selected JTS as a possible candi-
date for subducted sediment ( Fig. 16 ). The Sr and Nd
isotope composition of most post-16 Ma moredepleted back-arc basalts can be explained by the add-ition of <2% bulk sediment, and the most enriched iso-
tope ratios (0 /C17055–0 /C17065 for Sr
Iand 0 /C15125–0 /C15126 for
NdIat 15 Ma) of the subcontinental lithosphere can be
accounted for by addition of a maximum of 5–7% JTS,
if the original Sr and Nd compositions of the litho-
sphere approximated that of DMM. The Sr and Nd iso-tope composition of the frontal-arc basalts can beaccounted for by the addition of 1–5% JTS ( Fig. 16 b).
We have also examined modeling of partial meltingof M-DMM, using Nb/Yb versus La/Yb and Th/Yb
(Supplementary Data Electronic Appendix Figs A1
and A2).Implications for the generation of basalt magma
and tectonic setting during Japan Sea back-arcopening
From the discussion above, we propose a model to ac-
count for the tectonic setting and generation of basalt
magmas in the NE Japan arc during the pre-, syn- and
post-spreading stages of the Japan Sea back-arc basin(Fig. 17 ), which is summarized as follows.
1. During the early Oligocene (35–32 Ma) hot astheno-
spheric mantle (M-DMM) with a relatively depletedSr and Nd isotope signature began to upwell. Most
of the isotopically more enriched basalts of the back-
arc region were formed by melting of astheno-spheric mantle and subsequent interaction of the
melts with enriched lithospheric mantle or direct
lithospheric mantle melting. Slab melting occurred
and produced adakitic andesite at 34 /C14M a o n
Okushiri Island ( Sato et al. , 2014 ).
2. Asthenospheric upwelling continued from the early
Oligocene to the early Miocene (24–18 Ma) associ-
ated with the opening of the Japan Sea, and could
have reached the upper mantle beneath the frontal-arc region by about 20 Ma. More enriched basalts in
both the back-arc and frontal arc were derived from
asthenospheric mantle melting and subsequentinteraction with the subcontinental lithospheric man-
tle or by lithospheric mantle melting. Slab melting
also occurred at Okushiri Island, resulting in the
eruption of adakitic andesite similar to that formed
on the same island at 34 /C14M a( Sato et al. , 2013 ).
3. During the middle Miocene (16–12 Ma), the upwell-
ing asthenospheric mantle induced high-tempera-
ture conditions in the entire mantle wedge,
triggering slab melting to form the adakite magmasof the Ryozen area in the frontal arc ( Yamamoto &
Hoang, 2009 ). Most basalt magmas erupted in the
back-arc region owing to melting of upwelling as-thenospheric mantle had minimal interaction with
subcontinental lithospheric mantle, resulting in a
more depleted isotopic signature for these basalts.
Basalts of the frontal-arc region show more enriched
isotopic compositions owing to significant inter-action between the asthenosphere-derived magmas
and lithospheric mantle.
4. After 16–12 Ma, back-arc spreading ceased, astheno-
spheric upwelling weakened, and slab dehydrationbecame the main process for magma generation.The source mantle for the back-arc basalts at10–8 Ma, 6–3 Ma and 2 /C15–0 Ma may be the cooled
remnant asthenospheric mantle from which basaltmagmas were previously derived. The resultant
magmas interacted less with the overlying thin litho-
spheric mantle, resulting in isotopically moredepleted signatures for these basalts. Simultaneousfrontal-arc basalt magmas might have been gener-ated from a compositionally similar remnant as-thenospheric mantle source and subsequently
reacted with the overlying thinned lithospheric2288 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
mantle, resulting in a more enriched Sr and Nd iso-
topic character for these basalts. Mass balance ofsediment melt in the mantle and the process of man-tle melting have remained roughly constant sincethe cessation of Japan Sea opening (10–8 Ma, 6–
3 Ma and 2 /C15–0 Ma). Both the amount of sediment
melt added to the mantle wedge and the degrees ofmelting of the source mantle are higher for frontal-
arc magmas than for back-arc magmas.
CONCLUSIONS
To investigate the nature and origin of across-arc geo-chemical variations over time in mantle wedge-derivedFig. 17. Schematic cross-section of the NE Japan arc, showing the tectonic setting and production of basalt magmas during the
pre-, syn- and post-spreading stages of the Japan Sea back-arc basin. The NEJ basalt magmas at 35–0 Ma have been producedfrom three source materials: (1) a nearly homogeneous DMM-like mantle; (2) a source slightly more depleted than DMM; (3) the en-riched subcontinental lithospheric mantle. Prior to generation of these basalt magmas, such mantle sources have been modified by
subducting sediment-derived melt. High degrees of melting at shallow depths of 30–50 km in the mantle wedge beneath the frontal
arc persisted from 22 Ma to the present. In the back-arc, lower degrees of partial melting from 35 Ma to the present took place atdeeper levels (60–75 km depth) except for the middle Miocene (16–12 Ma) back-arc Tappizaki magma, which resulted from similardegrees of melting to the frontal-arc magmas during the same period. The model of material transportation (slab dehydration, hy-
dration of peridotite and decomposition of amphibole, chlorite and phlogopite) in a subduction zone is from Hanyu et al. (2006) .A
detailed description is given in the text.Journal of Petrology , 2015, Vol. 56, No. 11 2289Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
magmas, we have carried out a geochemical study of
basaltic rocks from both the frontal-arc and back-arc re-gions of the NE Japan arc, where basaltic eruptionshave continued intermittently from the early Oligocene(35 Ma) to the present.
The most fundamental aspects of the across-arc geo-
chemical trends for early Oligocene to Quaternary bas-alts in the NE Japan arc can be summarized as follows:(1) back-arc basalts at 24–18 Ma, 10–8 Ma, 6–3 Ma and2/C15–0 Ma show higher concentrations of HFSE such as
Nb and Zr, and have higher incompatible element ratios(e.g. Nb/Yb, Zr/Y, Zr/Lu and La/Yb) than coeval frontal-
arc basalts; (2) back-arc basalts have generally higher
REE (especially LREE and MREE) contents than the fron-tal-arc basalts, and hence the back-arc basalts exhibitsteeper chondrite-normalized REE patterns; (3) 35–32 Ma back-arc basalts (for which there are no coevalfrontal-arc basalts) exhibit similar HFSE abundances, in-
compatible element ratios, and REE abundances and
patterns to 24–18 Ma back-arc basalts; basalt magmaserupted at 16–12 Ma in both the frontal-arc and back-arcregions are usually characterized by low Nb/Yb and Zr/Yb, and high Nb/Yb and Zr/Yb, respectively. However,there are exceptions, including basalts from the
Tappizaki area, which were erupted in a back-arc region
but have trace element ratios similar to frontal-arc bas-alts. In general, melt-immobile mantle-derived HFSEconcentrations and incompatible element ratios in-crease from frontal arc to back-arc, from levels belowthose of DMM to levels approaching those of OIB.
Geochemical modeling of Nb/Yb variations versus
Nb shows that the frontal-arc and back-arc compos-
itional differences are independent of subduction modi-fication and can, in many cases, be explained bydifferent degrees of melting (higher degrees of meltingfor frontal-arc magmas and lower degrees of meltingfor back-arc magmas) of a nearly homogeneous DMM-
like source. However, some Pliocene frontal-arc basalts
may result from a source that is slightly more depletedthan DMM. Several 35–32 Ma and 24–18 Ma back-arcbasalts appear to be derived from lithospheric mantle,which is enriched in HFSE compared with DMM. A rare16–12 Ma basalt, which was erupted in the back-arc re-gion, was produced by a similar degree of melting to
frontal-arc basalts erupted at the same time.
Variations in ratios of fluid-mobile and -immobile
elements and those of melt- mobile and -immobileelements for the 35–0 Ma NE Japan basalts indicate thatthe principal subduction component added to thesource mantle prior to the generation of these basalt
magmas is a sediment-derived melt.
In terms of Sr
Iand Nd I, Pacific Ocean MORB, back-
arc basalts younger than 16 Ma and those older than18 Ma form a broad array extending toward JTS (bulksediment from the Japan Trench). The frontal-arc bas-alts from 24 Ma to the present also form a similar array.This suggests that the Sr and Nd isotope composition
of most post-16 Ma more depleted back-arc basalts can
be explained by the addition of <2% bulk sediment, andthe most enriched isotope values (0 /C17055–0 /C17065 for Sr
I
and 0 /C15125–0 /C15126 for Nd Iat 15 Ma) of the subcontinen-
tal lithosphere can be accounted for by addition of amaximum of 5–7% JTS, if the original Sr and Nd isotopecompositions of the lithosphere approximated that of
DMM. The Sr and Nd isotope composition of the fron-
tal-arc basalts can be accounted for by the addition of1–5% JTS.
The results of model calculations for partial melts of
DMM–sediment melt mixtures show that the Nb/Yb–La/Yb and Nb/Yb–Th/Yb variations of the 35–0 Ma NEJ bas-alts, excluding some Pliocene frontal-arc basalts and sev-
eral 35–32 Ma and 24–18 Ma back-arc basalts, can be
explained by different degrees of melting of DMM mixedwith<5% sediment melt. The isotopically more depleted
back-arc basalts (at 10–8 Ma, 6–3 Ma and 2 /C15–0 Ma) have
generally higher La/Yb and Th/Yb ratios than more en-riched frontal-arc basalts in the same periods. This may
be largely due to the more differentiated nature and ef-
fects of crustal assimilation for the former basalts.
We propose a depleted asthenospheric mantle
(DMM-like) upwelling model with interaction betweenasthenospheric mantle-derived magmas and the overly-ing lithospheric mantle to explain the geochemical char-
acteristics of the 35–0 Ma NEJ basalts. The frontal-arc
magmas were generally generated by higher degreesof melting of the shallower part of the asthenosphericmantle, whereas the back-arc magmas resulted fromlower degrees of partial melting of the deeper part ofthe asthenospheric mantle. These magmas wouldundergo interaction with the overlying lithospheric
mantle, resulting in more enriched Sr and Nd isotopic
signatures for pre-18 Ma back-arc basalts and post-22 Ma frontal-arc basalts, but less interaction, resultingin more depleted Sr and Nd isotopic signatures, formost of the back-arc basalts younger than 16 Ma.
FUNDING
This study was supported in part by the JSPS (theJapan Society for the Promotion of Science) Grant-inAid for Scientific Research (C) 25400510 (Rep. Prof.Kenji Shuto).
SUPPLEMENTARY DATA
Supplementary data for this paper are available at
Journal of Petrology online.
ACKNOWLEDGEMENTS
We would like to thank Toshiaki Takimoto for providing
samples from the Tomari area, Masahiko Suzuki forsamples from Sannome-gata volcano, Shin-ichi Tamurafor samples from the Joge and Mitaki areas, Jun’ichiOhki for samples from the Aosawa area, KazuhiroYamamoto for samples from the Fukuyama area, Hiroki
Yanagi for samples from Asakusa volcano, Seiko
Matsunaga for samples from Tobishima Island,2290 Journal of Petrology , 2015, Vol. 56, No. 11Downloaded from https://academic.oup.com/petrology/article/56/11/2257/2375424 by Ohio State University Law user on 12 March 2025
Masayuki Fukase for samples from the Kamo area,
Yumi Ogura for samples from the Myojin-iwa area.Thanks are due to Paul Morris for reading and improv-ing the paper. Critical and constructive reviews by JimGill, Christoph Beier, Marcel Regelous and the editorSimon Turner also improved the quality of this paper,and are greatly appreciated.
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www.nature.com/scientificreports scientific reports OPEN Feasibility o aroundwater dating calib both 81Kr and *He for the assessment of deep geological repositories in Japan 8 t'sn7 ue!1-buauz's'sbuey u!w ong ''sbue!r !m'Temebaseh ewnye1 'z'texyo oyowol Yasunori Mahara5 lodine-129, which is a promising tracer for dating old groundwater, has been used as a tracer for deep upwelling groundwater. The nuclide is expected to be one of the key factors for site selection for ratio for marine iodine is (1.50±0.15)x10-12, which could be considered the initial value for 1291 dating. This study identifies the challenges in groundwater age dating using 1291/1271. We measured the ratios of 1291/1271 and 81Kr/Kr and concentration of *He in groundwater from boreholes on the northern coast of Japan.The 129I dating results were not coincident with the other groundwater dating results.The iodine in the groundwater was inferred to be released in situ from marine organisms in sediments of various ages. We estimated that the primordial iodine ratio originating from seawater was ~ 1 x10-13 Ym saaibe anien s!ua buisn oes leztllez aua woy pasnpap ahe asempunob aui '(e-otxz~ st-otxg) other groundwater dating results. Limited groundwater mobility is one of the most important criteria for the safe and long-term disposal of high- level radioactive wastes in deep geological formations. Groundwater flow is the most critical factor in this assess- ment. Such assessments are important in Japan, where deep underground environments are rich in groundwater. An excellent way to determine groundwater mobility is to compare the residence time of the groundwater with the strata depositional age of the basement rocks that formed in the basin1-3. Most of the traditional meth- (T1/2 = 229 ka) has been used to determine the age of groundwater up to 1.3 Ma4.7-1 In addition, the presence of groundwater rising to the surface from deep underground indicates that a par ticular location is not appropriate for a repository site. This groundwater upwelling provides a pathway for radionuclides buried deep underground to move to the surface, where the nuclides could then contaminate human residential areas. This upwelling phenomenon is easily understood in cases where groundwater rises to the surface from deep underground and is visible at the surface. However, it is not clear whether groundwater Naturally occurring radionuclides in groundwater (*He, 81Kr, 36Cl) may be able to show whether the ground- water is older than the sedimentary strata deposited in a basin; therefore, these radionuclides can be used to determine whether the water that well up toward the ground surface at a potential disposal site is groundwater with a long residence time. When 81Kr is not detected and *He accumulation is extremely high in groundwater from great depth. Dissolved 1291 in groundwater is a long-lived cosmogenic radionuclide (T1/2 = 15.7 Ma), and the 1291/1271 ratio in groundwater has been used to estimate the residence time of groundwater with ages of tens of Ma12-14. Due to its long half-life, 129I cannot be used as a tracer for groundwater with ages lessthan 5 Ma, while 81Kr, *He and i291, it might be possible to verify the presence of groundwater with ages of tens of ka to 100 Ma. 1Civil Engineering Research Laboratory (Sustainable System Research Laboratory), Central Research Institute of Electric Power Industry, Abiko, Chiba 270-1176, Japan. 2Department of Nuclear Technology, Nagaoka University of Technology, Kamitomioka, Nagaoka, Nigata 940-2188, Japan. 3CAS Center for Excellence in Quantum Information and Quantum Physics, School of Physical Sciences, University of Science and Technology of China, Hefei 230026, China. *Hefei National Laboratory, University of Science and Technology of China, Hefei 230088, China. 5Kyoto University, Sakyo-Ku, Kyoto 606-8501, Japan.email: tomoohta@vos.nagaokaut.ac.jp Scientific Reports| (2024) 14:15688 |https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ gas field sitel2, the Horonobe sitel3, and the Kyushu sitel4 in Japan. In another study in Japan, the residence time of groundwater in the Kanto gas field estimated by the 36Cl and 4He dating methods coincided with the strata depositional age of the sedimentary rock where groundwater was stored?. On the other hand, the residence time estimated by the 1291 dating method significantly differed from the dating results2 based on the 36Cl and 4He concentrations. The pre-anthropogenic 1291/1271 ratio in marine iodine has typically been reported to be (1.50 ± 0.15) × 10-12, which could be considered the initial input value for 1291 groundwater dating15,16. Informa- tion on past ariations in isotopic ratios in seawater is stored in coral7.1 and algae samples5.1920oran et al.6 ratio in seawater estimated from pre-World War IIalgae was 1.5x 10-2 based on a mean of multiple algae samples collected from the pre-thermonuclear test eral5. The 1291/1271 isotopic ratio in coral was observed to be on the order of 10-2before194521Onthe other hand,the 1291/127isotopic ratiosofalgae samplescollected beforeWord War II and stored in museums have been observed tobe onthe order of 10-13 or more15,19 The observed 1291/1271 isotopic ratios in samples are widely divergent. The typical 1291/1271 isotopic ratios of natural sources are between 10-12 (Kilius et al.22) and 6 x 10-13 (Fabryka-Martin et al.23). Previous studies based on 36Cl (Mahara et al.1) and 4He (Mahara et al.2) in groundwaterl predicted that the primordial 1291/1271 ratio in groundwater ranges from 1.4 × 10-13 to 6 × 10-13 based on fossil seawater collected around the Kanto gas field in Japan. and meteoric water in coastal sedimentary areas to address the societal challenge of site selection for geological disposal. We proposed the primordial 1291/1271 ratio using the following methods: (1) an absolute groundwater dating method using dissolved 81Kr concentrations compiled from other dating methods based on 3H and 14C radioactivity and the dissolved noble gas concentration in deep groundwater collected from boreholes drilled in northern coastal Japan and (2) a new method focusing on biogenic elements (B, Br and C (TOC)) in seawater. The study site is defined as a sedimentary coastal area that sank into the deep sea in the pre-early Pleistocene (Fig. 1 and Fig. S1). Results Groundwater recharge age estimated from the radioactivity of 81Kr and 14C and the dissolved 4He concentration The study site was estimated to be subseafloor in the past, and coastal sediments continued to be deposited on the sea floor after the deposited sediments finally emerged from the sea surface (see Methods and Fig. S1). Supplementary Fig. S2 shows the major ions in the groundwater, which indicate that freshwater was distributed between the alluvium and Sarabetsu Formations and that brine included the freshwater distributed beneath the Yuchi Formation in the mixing zone. Based on the vertical distribution of the Cl concentration in the groundwa- ter, meteoric water infiltrated the Upper Yuchi Formation, and the original seawater was replaced with recharged Borehole : DD-1 141.69E Surface (0m) DD-4 DD-2 Sereen depth (GL-m) 214-215 N 306-307 476-477 DD-1 613-614 Study site 715-716 (Hamasato) 943-944 45.00N 1143-1144 Route 106 10m Well depth:1200m Japan sea Borehole : DD-2 Surface (0m) Screen depth (GL-m) okm 90.7-99.7 Well depth: 10 m Holocene 口 Alluvium Study Site Pleistocene 图 Terrace deposits Borehole : DD-4 Hamasato Sarabetsu Formation Surface (0m) Yuchi Formation 口 Koetoi Formation Sereen depth (GL-m) Wakkanai Formation 10km 337-348 Masuporo Formation Well depth:360m (a) (b) (c) Figure 1. Sampling site and geological map. (a) Geospatial Information Authority of Japan (GIS), (b) geological map and geological cross section of the site from Ikawa et al. 39, (c) location of three boreholes (DD-1, DD-2, DD-4) and screen depth of the boreholes. The edited map was obtained from the original map (https://www.gsi. Scientific Reports | (2024)14:15688| https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ at depths ranging from 90 to 1143 m, indicating that the recharge age of the groundwater in each formation was greater than 60 a (Table S1). Table 1 lists the concentrations of 81Kr and 85Kr in the groundwater and the cor- rected *1Kr concentration (pMKr, percent modern Kr). The corrected concentration of 81Kr in the groundwater ranged from 39 to 97.6 pMKr at depths ranging from 90.7 m to 99.7 m in the Upper Sarabetsu Formation and below 306 m between the Lower Sarabetsu and Lower Yuchi Formations. The 81Kr in the groundwater at depths the concentrations of Cl and s1Kr dissolved in the groundwater and the exchange period of seawater with mete- oric water deduced from the 81Kr dating results. The concentration of 81Kr in the groundwater in the Sarabetsu Measurementvalue s1Kr age recharge age Borehole Age upper 4He** cc ka limit (90% STP g1 Sedimentary Sampling 85Kr dpm 81Kr/81Kr_air confidence (recharge Geology age Ma Depth m date ccKr-1 81Kr pMKr (pMKr) err+ err- level) ka age) Alluvium ~0.02 0 90 90.7 99.7 DD-1 Oct.,3, 5.94±0.41 94.3 ±2.5 93.81 一 2.72 22.3 9.8 - 9.9 <10.8 2017 Upper 214 215 DD-2 Oct.,23, 0.9±0.3 104±4 104 4.05 13.5 13.5 - 13.5 <13.1 Sarabetsu F 2017 6-10× 10-8 1.3* (0.018- 214 215 DD-2 Oct.,23, 1.0±0.2 103±4 103 士 4.05 10.1 13.5 - 13.4 <28.4 0.034 Ma) 2017 Lower 306 307 DD-1 Jun. 7,2017 3.19 ±0.41 97.7 ±3.4 97.6 3.55 8.5 13 -12.3 Sarabetsu 613 614 DD-1 Sep.,13, 1.7 ± 0.3 88±3 87.7 3.1 44.1 12.0 -11.0 Upper 2017 Yuchi F DD-1 May 21, 715 716 33.2±3.9 72.0±5.8 49.8 11.4 231 86 -68 2017 Nov., 16, DD-1 38.3±1.3 2-9×10-6 2 943 944 75±4 48.9 ± 8.4 237 62 -52 一 2017 (1-2 Ma) Lower 943 944 DD-1 Nov., 16, 38.1±2.5 Yuchi F 70±6 39.0 12.9 312 137 - 94 2017 Nov., 17, 943 944 DD-1 2017 NA NA Table 1. Concentrations of 8Kr,8Kr, and *He in the groundwater and estimated recharge age of the groundwater. *Ikawa et al.39 **in-situ sampling from 214 to 944m, GL sampling above 200m (Hasegawa et al.24) 8 groundwater 0 0 0 Alluvium Upper -200 -200 -200 口 Sarabetsu Meteoric water Sarabetsu 口 F. Lower -400 -400 -400 Sarabetsu Mixing zone (m) (m) Depth ( -600 Depth Yuchi F. -600 Depth 600 口 8 口 -800- -800 -800 0口 -1000 -1000 -1000 Fossil seawater -1200- -1200- 1200- 0 10000 20000 20 40 60 80 100 120 -1000 100200300400500 Cl (mg L-) 81Kr/81Krair (pMKr) 81Kr (ka) Figure 2. Concentration depth profiles of Cl, 81Kr, and estimated recharge age of groundwater. O: concentration of Cl (red circle: groundwater, black circle: pore water), O: concentration of s1Kr, : recharge age of groundwater estimated from 81Kr Scientific Reports | (2024)14:15688 https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ Formation was 94-98 pMKr from a depth of 90 m to 307 m, corresponding to a residence time of 9-22 ka. The meteoric water is less than the estimated 22 ka based on 81Kr dating of the groundwater. The concentrations of 14C and 4He in the groundwater from a depth of 90-307 m in the borehole ranged from 15 to 22 pMC and from 6× 10-8 to 10 × 10-8 cc STP g1 (STP, standard temperature pressure), respectively (Tables 1 and S1). The estimated groundwater ages at the same depth ranged from 8 to 13 ka based on 14C and from 18 to 34 ka based on *He24. The recharge era of these meteoric waters estimated by the 81Kr dating results was roughly in line with the ages estimated from both the dissolved concentrations of i4C and 4He in the Sarabetsu Formation. The results esti- mated from the 81Kr, 14C and dissolved 4He concentration dating methods clearly showed that the groundwater found in the Sarabetsu Formation originated from meteoric water discharged approximately 22 ± 10 ka. The concentration of 81Kr dissolved in the groundwater in the Upper Yuchi Formation, corresponding to the mixing zone between freshwater and brine (connected with seawater), was lower than that in the Lower Sarabetsu Formation. The 81Kr concentration in the Upper Yuichi Formation decreased with increasing concentrations of Cl and Br, and the 8D value also increased with increasing 818O value. Finally, the estimated 81Kr age increased with increasing groundwater sampling depth in the formation. Although the Cl concentrations in the groundwater almost constant at 0.16-0.45 Ma. Assuming that seawater was trapped as pore water before 2 Ma when ancient seawater permeated the Yuchi Formation (Fig. S1), the detection of8Kr dissolved in the seawater that intruded into rock pores can be diffcult as radioactive decay proceeded for 2 Ma. The presence of groundwater with a Cl concentration (17,0o0 mg L-l) lower than that of seawater could be due to groundwater dilution by water released from the dehydration of organic matter and minerals25. Alternatively, if brackish water from lagoons and/or rivers permeated into the Yuchi Formation approximately 0.15-1 Ma, this permeation might explain the lower Cl concentration relative to that of modern seawater (e.g., the salinity of Tokyo Bay is 22-34%o with an Yuchi Formations suggested that the age of the seawater and/or brackish water from the shallow bay and lagoon that intruded into the Yuchi Formations was at least 0.31 Ma (at most 0.45 Ma). Dissolved 4He in the ground- water was collcted by in situ sampling (Fig. S3b,c). Unlike typical sampling methods based on groundwater pumping, this method prevents the sampled groundwater from degassing and protects it from contamination by the modern atmosphere. The concentration of 4He dissolved in the groundwater collected in situ from the Yuchi Formation was on the order of 10-6 cc STP g-1 (2-9 ×x 10-6 c STP g-'), which corresponded to mud pores. where 4He was produced in situ and accumulated in the host sedimentary rock. The accumulated amount of 4He in the groundwater indicates an accumulation time of 1-2 Ma, which is the estimated residence time of the groundwater24, and the age estimated from this 4He accumulation rate roughly corresponds to the geological sedimentary age of the transformation from a deep bay environment to a shallow lagoon environment. The groundwater in the Yuchi Formation originated from fossil seawater, the brackish water environment changed from a deep bay environment to a shallow lagoon environment in less than 2 Ma, and meteoric water intruded before 0.04 Ma and was found at 614 m in the formation. Groundwater recharge age estimated from the 1291/1271 ratio in groundwater s0x o T v8 ot 6o uo n Iio son sdoosi e I pe L Jo suousu u to 7 ×107 atoms L, and from 6 × 10-14 to 17× 10-14, respectively (Table 2). Figure 3 shows the concentration e It pnpap u pue (oner sruoe) Iz/ pue (T suoe) Ii (T suoe) II Jo syod ydap snsia of the groundwater estimated from the traditionally accepted initial ratio of 1291/1271 (1.5 × 10-12). The 1291/271 ratios in the groundwater were generally on the order of 10-13. The estimated 129I age based on the initial ratio of s m o s o s o from 3 tracer methods: the 81Kr, *He and traditional 1291 groundwater dating methods. The estimated recharge strata depositional age at the observation site. We calculated the recharge age of groundwater for a given initial 1271 1271 1291 1291/127 ×1020 x106 X10-14 Geology Depth m Borehole mg L-1 atoms L-1 atoms L-1 SD SD 90.7 99.7 DD-2 9.6 0.5 5.38 0.86 11.8 1.9 Upper Sarabetsu F 214 215 DD-1 19.0 0.9 7.91 1.23 8.8 1.4 306 307 DD-1 1.2 0.1 0.388 0.118 6.55 1.99 LowerSarabetsuF 337 348 DD-4 0.9 0.0 0.600 0.100 14.8 3.0 476 477 DD-1 48.3 2.3 32.6 1.39 14.2 0.6 613 614 DD-1 66.0 21.7 4.00 6.9 Upper Yuchi F 3.1 715 716 DD-1 84.3 4.0 23.9 3.40 6.0 0.9 943 944 DD-1 66.9 3.2 37.32 2.87 11.8 0.9 Lower Yuchi F 1143 1144 DD-1 84.3 4.0 68.0 4.0 16.5 0.5 Table 2. Concentrations of 1271 and 1291 and the 1291/1271 ratio in the groundwater. ScientificReports| (2024)14:15688| https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ 0 0 O 口 Alluvium 一 Sarabetsu F. -200 O Meteoric -200- 200 -200 口 Upper Sarabetsu water ， 口口 Lower Sarabetsu -400 -400 -400 -400 O 口 Mixing Depth -600 zone Depth O -600 Depth 600 Depth 600 口 O 口 Yuchi F. -800 Fossil -800 -800 -800 seawater O 口 -1000 -1000 -1000 -1000 O 口 -1200 -1200- -1200 traditional -1200 inital ratio 0 5 10 0 50 100 0 50 100 150 0 50 100 1271 (x1020 atoms L-1) 1291 (x10° atoms L-) 1291/1271 (×10-14) 1291 age (Ma) groundwater based on the traditional 1291 groundwater dating method. O: concentration of 127 in groundwater, O: concentration of 129I in groundwater, : 1291/1271 ratio in groundwater, : recharge age of groundwater estimated from traditional i291 dating, ---:traditional initial ratio (1.5× 10-12 by Fehn et al. 15). s1Kr age Sedimentary 4He age Geology ageMa Depth m Ma err- Ma initial ratio 1.5 × 10-12 Ma 1.65 x 10-13 Ma 1.18 x 10-13 Ma 8.80 x 10-14 Ma Alluvium ~0.02 0 90 一 一 一 一 90.7 99.7 58 8 0 -7* Upper Sara- 214 215 0.022 0.010 -0.010 64 14 7 betsu F 0 214 215 0.018-0.034 306 307 Lower Sara- 71 21 13 betsu 1.3 337 348 52 2 -5* -12* 476 477 53 3 -4* -11* Upper Yuchi F 613 614 0.04 0.012 - 0.01 70 20 12 5 [23 715 716 0.23 0.086 - 0.07 73 15 9 943 944 0.24 0.062 -0.05 1-2 943 944 0.31 0.137 - 0.09 0 *L- Lower Yuchi F 58 943 944 一 1143 一 1144 一 -8* 14* Table 3. Comparison of the residence times of groundwater samples determined from 3 tracer methods: the 81Kr,*He and traditional 1291 groundwater dating methods. *Negative values indicate that the initial value is smaller than the isotopic ratio in groundwater. **The 1291/1271 ratio in groundwater with recharge ages up to 5 Ma is within the error range of the initial values. value of 1291/1271 (1.5 × 10-12~8.80 × 10-14). These groundwater ages were 0~23 Ma for 1.65 × 10-13, - 8~15 Ma for 1.18× 10-13 and - 14~9 Ma for 8.80 × 10-14. These age also did not agree with groundwater age estimated from other groundwater ages. lodine and biogenic elements Figure S4 shows the concentrations of 1271 and 1291 in the groundwater with respect to TOC content. The 1271 con- centration was correlated with TOC (correlation coefficient (R?) =0.819, p-value (P) = 0.000797<0.05). However, the 129I concentration was not correlated with the TOC content (R2 = 0.4924, P= 0.0847 >0.05). The I/Br ratio (mol/mol) in the groundwater ranged from 0.2 to 3.6 at the site (Fig. S5). The concentration of Scientific Reports | (2024)14:15688 https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ increased with increasing boron content, while it did not increase in other samples (Fig. S6b). The 1291/1271 ratio showed no correlation with the boron concentration (R2 = 0.0278, P=0.67 >0.05) (Fig. S6c). We normalized 1291/1271 to the concentration of TOC (mg L-1) to determine the effect of the addition of old organic matter with low 1291/127I values. The ratio of (1291/i271)/TOC decreased with increasing TOC concentra- tion (Fig. 4). The regression curve between (1291/1271)/TOC and TOC was (1291/1271)/TOC = 8.404x (TOC) -0.956 (R2 =0.9367), as shown in Fig. 4. We also normalized 1291/1271 to the concentration of boron (mg L-1). The ratio of (1291/1271)/B in the groundwater decreased with the concentration of boron (Fig. S6d). The regression curve between boron and (1291/1271)/B was (1291/1271)/B = 10.1 x (B) -0.995, as shown in Fig. S6d (R2 = 0.9393). Discussion Assuming that the groundwater residence time estimated from the present 1291 dating method (50-71 Ma) is correct, the groundwater has risen from a deep underground formation to the ground surface. If the sedimen- tation process at this site was complete between the Pleistocene and the Pliocene, the estimated groundwater age exhibited large discrepancies between the two prediction methods (groundwater dating by 81Kr and +He vs. Sarabetsu and Yuchi Formations is not older than that in the strata based on our groundwater dating results. This contradiction among the estimated groundwater ages cannot be resolved, even if the traditionally accepted a ( h ss s su (e a) e on st (ioi xs'r) Ieian ro o groundwater at the site does not well up from deep underground and 2) other factors must be considered to determine the recharge age of groundwater estimated by 129]/1271. The iodine dissolved in the groundwater in the Sarabetsu Formation originated from only two sources; the main source was meteoric water that recharged in the last ~ 40 ka, and the other source was dissolution from organic matter in the sedimentary layers that settled offshore between 1 Ma and 0.018 Ma. The iodine in the groundwater in the Yuchi Formation has three origins: (1) seawater and lagoon water from when the coastal sediment was deposited, (2) infiltrated meteoric water from before 0.04 Ma (corresponding to the 81Kr age in groundwater at the Upper Yuchi Formation in Table 1), and (3) iodine released from organic matter in the sedi- ment. The concentrations of 1271 supplied from meteoric water and seawater are only several μg L (Yuita et al.28) and ~ 58 μg L-1 (Elderfield and Truesdale29), respectively. Iodine is easily adsorbed on sediments30,31 and tends to accumulate in organic matter32, cells of marine diatoms33 and humic substances34. Not all the iodine dissolved into the pore water of the formation; some remained in the sediments as organic iodine during the maturation of organic matter during diagenesis13 Considering that the concentration of 1271 in the groundwater ranged from 10 to 19 mg L-1 in the Upper Sarabetsu Formation and from 66 to 84 mg L-l in the Yuchi Formation, the main source of iodine dissolved in the groundwater would be iodine released from organic matter in the sediments. The organic matter deposited in the sediment has various origins: (1) coeval ocean and (2) old recycled organic matter. We hypothesized that (1) iodine was released into the groundwater through the decomposition 1000 (1291/1271)/TOC ■ seawater (traditional) seawater (this work) 100 y=8.404x (-0.956) L w L (×10-14/ ■ y = 8.4037x-0.956 10 R2=0.9367 0.8 ■ 0 '1)/TOC 0.6 OC 0.4 127 /71 0.2 0.1 12 0.0- 0 50100 150 200 250 300 350 0.01 TOC (mg L-l) 0.1 10 100 1000 TOC (mg L-) (b) (a) Figure 4. Correlations between the (1291/1271)/TOC ratio and the concentration of TOC (mg L-1) in : seawater (Traditional): (1291/1271)/TOC values of 179 × 10-14 and 224× 10-14 calculated from 1291/1271 (150 × 10-14 traditional initial ratio estimated by Fehn et al. 15) and TOC concentrations in surface seawater of 0.67 and 0.84 mg L-,respectivly38 : seawater (This study): (1291/271)/TOC values of 12.3× 10-14 and 9.93× 10-14 obtained from the regression curve of the groundwater values with TOC concentrations of 0.67 and 0.84 mg L-, respectively. Scientific Reports| (2024) 14:15688| https://doi.0rg/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ iodine isotopic ratios observed in the groundwater. As the biogenic ratio of I/Br in marine organic materials is greater than 0.04±0.02 (Elderfield and Trudesdale29, Pedersen and Price35),both the I and Br dissolved in the groundwater could be of biogenic origin. Surface seawater has a boron concentration of 4.5 mg L-1; however, boron can be efficiently concentrated into marine products3637 (more than 200 mg L-1). The iodine concentration in groundwater increases with increasing boron content, which is released via organic matter decomposition. ing from marine organisms (algae, marine plankton) in the sediments. The relationship between the ratio of (1291/1271)/TOC and TOC shows the effect of the addition of old organic matter with low i291/1271 values (Fig. 4). This sugests that the groundwater contains recycled organic matter that is older than the sedimentary age of the stratum. Recycled organic matter could be present in any stratum because some recycled organic matter settles on the seafloor due to inflow from river sediments and some is dissolved in seawater. Here, the primordial 1291/1271 ratio in groundwater is defined as the value when the original seawater is recharged to the surface of the sediment. As the iodine originated from marine products that became concen- trated from seawater, the primordial 1291/271 in the groundwater indicates the value in the sea and lagoon seawater deposited in the Yuchi Formation and Sarabetsu Formation. The concentration of TOC in surface seawater in the open ocean is ~ 58-70 μM (corresponding to a TOC of ~ 0.67-0.84 mg L-1) (Figure A2 in Appendix A by Halewood et al.38). Here, assuming that the TOC in surface seawater is 0.67-0.84 mg L-, the (1291/1271)/TOC e a a s xa) si o curve. However, if the TOC and 1291/1271 values in surface seawater are 0.67-0.84 mg L-1 and 150 x 10-14 (corre- sponding to the traditional ratios), respectively, the (1291/271)/TOC ratio will be (179-224)× 10-14 (Fig. 4). This is a significant departure from the extrapolated values from the regression curves between TOC and (i291/271)/ TOC obtained from the groundwater samples in Fig. 4. Here, when the concentration of boron in surface seawater is 4.5 mg L,the (1291/121)/B ratio estimated from the traditional 1291/271 model is 33× 10-14, which is plotted well outside the regression curve. However, based on the regression curve in Fig. S6d, when the concentration of boron in the surface seawater is 4.5 mg L-, that of (1291/21)/B is 2.26 × 10-14. Therefore, the 1291/1271 ratio in the surface seawater when the concentration of boron is 4.5 mg L-l may have been 10.2 x 10-14 (~ 1 x 10-13). The 1291/1271 value in surface seawater estimated from (1291/1271)/TOC and TOC is similar to that estimated from (1291/1271)/B and B. Therefore,the 1291/1271 ratio in seawater when the sediment was deposited between the Yuchi Formation and the Sarabetsu Formation was estimated to be ~ 1 x 10-13 based on the regression curve in Fig. 4. (2 Ma), the amount of 291 (atoms L-l) that decayed out of the total iodine dissolved in the groundwater in the Yuchi Formation to the Sarabetsu Formation is negligible. Therefore, the value of (16.5±0.5)x 10-14 (~2x 10-13) for the 1291/1271 ratio in the groundwater in the Lower Yuchi Formation includes the primordial ratio. Assuming that the primordial ratio is ~ 1 x 10-13 (8 × 10-14 ~ 2 × 10-13), the 1291/1271 ratio in seawater is almost the same as the primordial ratio. As the 1291/1271 ratio in the groundwater is almost in the same range of 10-13 m u oi m se aa o o od a sm m u o recharge ages up to 5 Ma within the margin of error. The age of groundwater originating from fosil seawater in the Pleistocene stratum in the Kanto gas field is less than 1 Ma, as estimated by 36Cl and *He groundwater dating2. The 1291/1271 ratio in groundwater originating from fossl seawater in the Kanto gas field is ~ 2 × 10-13. The ratio for the fossil seawater is similar to our estimated primordial ratio for the seawater. To obtain accurate groundwater ages in coastal sedimentary areas based on the 1291/1271 ratio in ground- water, the following values need to be clarified: the primordial ratio in the original groundwater recharged in the stratum, the 1291/1271 ratios in the endmembers of the old organic matter and the corresponding abundance ratios. Even the 1291/1271 ratio in recycled organic matter older than the sedimentary age of the stratum will not possible with 1291_129Xe in the future for groundwater with a residence time of tens to 100 Ma. Conclusions 1.5 ×10-12 was more than three orders of magnitude higher than the ages estimated from other (8Kr, 14C, and 4He) dating methods. The 1291/271 ratio in the groundwater ranged from 6×10-14 to 2 x 10-13, suggesting that iodine in organic matter of various ages was released into the groundwater in the stratum. We propose that the dsod o oxx ex s s ss u o d 1291/1271 ratio, all the groundwater ages are consistent (within 5 Ma). Methods Field site The boreholes at the investigation site are located in Hamasato on the coast of Hokkaido, Japan, as shown in Fig. 1. The drilling location was 300 m inland from the coastline at an elevation of approximately 5.2 m. Fig- ure 1b shows the geological cross-sectional map of the site. The detailed geological and porewater information for the survey site was described in a previous study39,40. In brief,the sediment at the observation point has the of shallow marine sediment, was deposited from a depth of 2 km to 471.5 m during a period from 2.3 Ma or Scientific Reports | (2024) 14:15688| https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ younger in the late Pliocene under an offshore environment. (2) The Sarabestu Formation has overlain the Yuchi Formation since 1.3 Ma in the early Pleistocene, and the site changed from a bay to a lagoon environment. (3) The Sarabestu Formation continued to be deposited in the period from 0.8 to 0.15 Ma (interglacial) in the Middle Pleistocene. The sedimentary environment of the site changed from a lagoon to a river environment. (4) Alluvium was deposited since approximately 10 ka at a thickness of 85 m. That is, at this site, meteoric water started to recharge vertically from the surface alluvium between 0.8 Ma and 0.15 Ma when the lagoon changed to a fresh- depths from the surfaceto the SarabetsuFormation, suggesting that meteoric water displaced the fossil seawater that had existed from the surface to the Sarabetsu Formation and penetrated since O.8 Ma from the alluvium to the Sarabetsu Formation (Fig. S1). Although fossil seawater remains in the Yuchi Formation, the pore water in the upper part of the Yuchi Formation is a mixture of meteoric water and fossil seawater. The groundwater in the present aluvium and the Sarabetsu Formation consists of fresh water, while the groundwater in the upper part of the Yuchi Formation is a mixture of meteoric and fossil seawater, and the Cl concentration in the Yuchi Formation (- 1200 m) is between 10,000 mg L-d and 17,000 mg L-. Sampling of deep groundwater at the site depths of 1200 m (DD-1 borehole), 100 m (DD-2 borehole) and 360 m (DD-4) (Fig. S3). The screen depths were 214-215 m, 306-307 m, 476-477 m, 613-614 m, 715-716 m, 943-944 m, and 1143-1144 m for the DD-1 borehole; 90.7-99.7 m for the DD-2 borehole; and 337-348 m for the DD-4 borehole, as shown in Fig. 1c. The groundwater samples were collcted within the perforation intervalusing an upper packer and lower packer from boreholes drilled to depths ranging from 90.7 to 99.7 m and 214 to 215 m for the Upper Sarabetsu Formation; from 306 to 307 m and 337 to 348 m for the Lower Sarabetsu Formation; from 476 to 477 m, 613 to 614 m, and 715 m to 716 m for the Upper Yuchi Formation; and from 943 to 944 m and 1143 to 1144 m for the Lower Yuchi Formation. Groundwater for radio-Kr, 14C, radioiodine, and other chemical analyses was pumped and sampled above ground (Fig S3a). After sampling, the noble gases present in the groundwater were collected into copper tubes in situ (Fig. S3b,c). The in situ sampler mainly consisted of a piston sampler, copper tubes (three vertically of groundwater. At the target depth, the pressure of the piston sampler was reduced, and the groundwater was passed through the copper tube. After sampling was complete, the piston sampler was pressurized to close the check valve and raised to the ground. The field campaign was held in 2017-2018. 1291 Groundwater samples were collected in plastic bottles. The method for the extraction and purification of iodine was as follows. The dissolved iodine was oxidized to I2 and was then separated from the sample water into 10 ml of organic solvent (CCl4 or hexane). The solvent was separated from the sample water, and then 10 ml of 0.1 M Na2SO or NaHSO3 was added to the solvent to extract I into the Na2SO3 or NaHSO, solution. The Na2SO3 or NaHSO, solution was separated from the solvent, and 0.1 ml of concentrated HNO, was added to the solution. A 0.1 ml portion of 1 M AgNO3 was then added to the solution, and the solution was agitated. Precipitation was separated from the solution. The AgI precipitate was rinsed with 5 ml of ultrapure water to yield a pure AgI sample, which was then dried in an electric oven at 80 °C for 12 h. The 1291/1271 isotopic ratio in the AgI sample was measured with an accelerator mass spectrometer (Ottawa University and The University of Tokyo). When Ia e sm oe sqon s sm o e isotopic ratio of the carrier for isotope dilution was 3± 1 × 10-14. The 1271 concentration in the groundwater was measured by ICP-MS. 81Kr the necessary samples for 81Kr dating?. Depending on the methane content of the dissolved gas, the extracted Kr values ranged from 0.2 to 10 μL STP. The relative isotopic abundances of both 81Kr and 85Kr were measured by atom trap trace analysis (ATTA)9. Since the half-life of 85Kr is only 10.7 a, any presence of 85Kr is an indication of modern air contamination, which could be introduced in wellheads, pumping systems or sample treatment processes. Assuming that the contamination is from local air during sampling, the 85Kr activity can be used to make contamination corrections to the measured 81Kr abundance (Rs1_meas). The 81Kr abundance (R81_cor) after correction can be expressed as follows: Rs1_corr = (R81_meas - 100 ×α)/(1 - α) (1) where α is the fraction of Kr from modern contaminants. The contamination fraction α can be determined using the measured 85Kr activity (Rs5_meas) and the 85Kr activity of the local air at the sampling site: α = R85_meas / (R85_atm) (2) BycombiningEqs. (1 and(2),thecorrected 8Krabdance (Rs_co) canbe calculate. The value of Rs am is assumed to be 75 dmp/cck 43-45 Scientific Reports| (2024)14:15688| https://doi.org/10.1038/s41598-024-66250-3 nature portfolio www.nature.com/scientificreports/ Analytical methods for the chemical characterization of groundwater We measured stable hydrogen and oxygen isotope ratios (8D and i8O), as well as the concentrations of eight major dissolved ions (Na*, K+, Ca2+, Mg2+, Cl-, SO42-, NO; ) and dissolved trace elements (total amount of Br, I, and B) in the samples. The major ion and trace element concentrations were measured by ion chromatography Agilent Technologies Inc., CA, USA), respectively. SD and 818O values were measured by the cavity ring-down method (IWA-35EP, isotope ratio mass spectrometry, Los Gatos Research, San Jose, CA, USA) and Iso-Prime (GV Instruments, Manchester, UK), respectively. The groundwater was distilled for 3H analysis and then subjected to electrolytic enrichment of tritium. The enriched tritium samples were mixed with a scintillation cocktail. The radioactivity of the tritium in the samples was measured by a liquid scintillation counter (LSC-LB5, Aloka, Co. Ltd.). Other groundwater dating data collcted at the borehole site were estimated from noble gases (He, Ne, Ar, Kr, and Xe) and 14C24. The 14CO2 in the groundwater was separated by gas stripping6 and then further purifed to obtain 14C. The 14C/C ratio was measured by accelerator mass spectrometry (AMS) (Beta Analytical Co. Ltd., USA). TOC was measured with a nondispersive infrared gas analyzer. 813C was measured by a stable isotope mass spectrometer (Beta Analytical Co. Ltd., USA). Data availability All data generated or analyzed during this study are included in this published article (and its Supplementary Information fles). Received: 30 October 2023; Accepted: 29 June 2024 Published online: 08 July 2024 References 1. 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Acknowledgements This study was conducted under a contract with the Ministry of Economy, Trade and Industry (METI) as part of the R&D support program tiled “Development of enhancing the disposal system in the coastal region (2019Fy)" (Grant Number JPJ007597). The original idea and basic research for determining the primordial ratio of 1291/1271 in groundwater were established in the period from 2008-2010 (supported by a Grant-in Aid for Sci- entific Research A, KAKENHI number 20241006), and basic research was partly supported by KAKENHI Grant Numbers 24686098 and 20H02674. This work has been supported by the Innovation Program for Quantum Sci- ence and Technology (2021ZD0303101), the National Natural Science Foundation of China (41727901). We and Matsumoto) for their cooperation and CRIEPI members (Dr. Tomioka and Mr. Okamoto) for groundwater sampling. Author contributions T.O. and Y.M. developed the original idea for the analyses presented in the manuscript. Field data were collected by T.O. and T.H. Laboratory analyses were performed by T.O., W.J., G-M.Y., Z-T.L., and T.H. The manuscript was written by T.O. with contributions from all coauthors. Competing interests The authors declare no competing interests. Additional information Supplementary Information The online version contains supplementary material available at https://doi.org/ 10.1038/s41598-024-66250-3. Correspondence and requests for materials should be addressed to T.O. Reprints and permissions information is available at www.nature.com/reprints. Publisher's note Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. Scientific Reports | (2024)14:15688| https://doi.org/10.1038/s41598-024-66250-3 nature portfolio 10 www.nature.com/scientificreports/ Open Access This article is licensed under a Creative Commons Attribution 4.0 International License, which permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons licence, and indicate if changes were made. The images or other third party material in this article are included in the article's Creative Commons licence, unless indicated otherwise in a credit line to the material. If material is not included in the article's Creative Commons licence and your intended use is not permitted by statutory regulation or exceeds the permitted use, you will need to obtain permission directly from the copyright holder. To view a copy of this licence, visit http://creativecommons.org/licenses/by/4.0/. The Author(s) 2024, corrected publication 2024 Scientific Reports| (2024)14:15688| https://doi.org/10.1038/s41598-024-66250-3 natureportfolio 11
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Tectonic shortening and coeval volcanism during the Quaternary, Northeast Japan arc KOJI UMEDA1*, MASAO BAN?, SHINTARO HAYASHI? and TOMOHIRO KUSANO1 1 Tono Geoscientifc Research Unit, Geological Isolation Research and Development Directorate, Japan Atomic Energy Agency, 959-31, Jorinji, Izumi, Toki, 509-5102, Japan. 2 Department of Earth and Environmental Sciences, Yamagata University, 1-4-12 Kojirakawa-machi, Yamagata, 990-8560, Japan. 3 Department of Natural and Environmental Sciences, Akita University, 1-1, Gakuen-machi, Tegata, Akita, 010-8502, Japan. *Corresponding author. e-mail: umeda.koji@jaea.go.jp The Northeast Japan arc, a mature volcanic arc with a back-arc marginal basin (Japan Sea), is located on a convergent plate boundary along the subducting Pacific plate and the overriding North American plate. From a compilation and analysis of stratigraphy, radiometric age and data on erupted magma volumes, 176 eruptive episodes identified from 69 volcanoes so far, indicate that notable changes in eruption style, magma discharge rates and distribution of eruptive centres occurred around 1.0 Ma. Before ca. 1.0 Ma, large-volume felsic eruptions were dominant, forming large calderas in the frontal arc, a region of low crustal strain rate. After ca. 1.0 Ma to the present, the calc-alkaline andesite magma eruptions in the frontal and rear arcs, synchronous with crustal shortening characterized by reverse faulting, resulted in stratovolcano development along narrow uplifted zones. Although, it is widely assumed that magma cannot rise easily in a compressional setting, some of the magma stored within basal sills could be extruded where N-S-trending uplifted mountains bounded by reverse faults formed since about ca. 1.0 Ma. 1. Introduction Watanabe et al. 1999). However, volcanism is espe- cially common in areas of extensional tectonics, At highly active compressive margins, such as and notably so at divergent plate boundaries (e.g., the Northeast (NE) Japan arc, current stresses Galland et al. 2007). Similarly, numerous Quater- are mostly compressive and the greatest principal nary volcanoes are distributed in the NE Japan stress is horizontal, indicating that the major faults arc, where the Pacific plate is subducting along the are thrust faults (e.g., Nakata and Imaizumi 2002). Japan Trench under the North American plate at In compressional settings, it is considered that a rate close to the modelled convergence of nearly magma cannot ascend easily, favouring magma 9 cm/yr (DeMets 1992). Indeed, the global posi- upwelling along vertical fractures perpendicular to tioning system measurements reveal large contrac- the regional least principal stress (o3) (Hubbert tion (>0.1 ppm/yr) under compression in the arc and Willis 1957). It has been suggested by many (e.g., Sagiya et al. 2000). This region is probably authors that volcanism should be rare if not absent one of the most extensively studied volcanic arcs in under horizontal compression (e.g., Hamilton 1995; the world, particularly regarding the relationship Keywords. Northeast Japan arc; volcanism; tectonic shortening; Quaternary. J. Earth Syst. Sci. 122, No. 1, February 2013, pp. 137-147 Indian Academy of Sciences 137 138 Koji Umeda et al. between volcanism and tectonics, and high-quality supplied from locally developed hot regions within geological, geophysical and geochemical data are the mantle wedge (Tamura et al. 2002). In the NE readily available (e.g., Hasegawa et al. 1978; Zhao Japan arc, a change in the stress field at around the et al. 1992; Sato 1994; Umeda et al. 1999; Yoshida Pliocene to Quaternary boundary resulted in E-W 2001; Tamura et al. 2002; Kimura and Yoshida compression (Sato 1994). In connection with the 2006; Acocella et al. 2008). tectonic sequence, felsic volcanism associated with It has recently been recognized that Quaternary the large calderas terminated and andesite strato- volcanoes are clustered in regions trending trans- volcanoes became the dominant form of volcanism verse to their host arc axis, and separated by sub- during the Quaternary (Sato 1994). A compressive parallel, volcano-free gaps 30-75 km wide along the stress regime is not favourable to the development arc (figure 1) (Umeda et al. 1999; Tamura et al. of felsic volcanism, requiring the formation of a 2002). The arrangement of the volcanic clusters, large magma reservoir at shallow depth (Yoshida coupled with topographic profiles and negative 2001). The predominance of stratovolcanoes can Bouguer anomalies, could be controlled by magma be reconciled with the crustal stress field in the present-day subduction system. In contrast, magma discharge rate is inferred to have increased in the NE Japan arc in the late Quaternary. A com- pilation of age data indicates that volcanoes rang- ing in age from 2.5 Ma to 1.5 Ma are relatively rare compared with those younger than 1.5 Ma (Kimura and Yoshida 2006). This insight appears to be con- tradictory to the idea that compressive stress in Japa the Quaternary would prevent the rise of magma within the crust. Recently, new stratigraphic and radiometric ages Pacific of Quaternary volcanoes in the NE Japan arc have been reported, and several synthesized databases were compiled by Umeda et al. (1999) and Com- 130° 135° 40° 145 mittee for Catalogue of Quaternary Volcanoes in Japan (1999). Therefore, the special and temporal variations in eruption style, magma discharge rate and distribution of volcanoes during the Quater- nary can be recognized with higher resolution than 41 before. With the merging of this type of informa- tion, an understanding of the geometric and kinetic features of the structures controlling magma rise and emplacement in convergent settings can be developed. This understanding is also important to 40- applied studies of natural hazards and economic resources. For this paper, we compiled the latest radiometric and stratigraphic data, and calculated the dense-rock equivalent (DRE)(Walker 1980）of volcanic products for each volcano. We describe time-magma volume behaviour that could provide insight into the tectono-magmatic relationships throughout the volcanic arc during the Quaternary. In addition, we examined the mechanism of magma transfer within the crust in a compressional setting. 02550 100km 2. Geological background 37 138 139 140 141 142 143 The Cenozoic volcanic and tectonic sequence in Figure 1. Location and shaded relief maps of NE Japan and the NE Japan arc is directly associated with the distribution of Quaternary volcanoes (black triangles) and separation of the present-day NE Japan arc from active faults (black lines) (Nakata and Imaizumi 2002). Solid triangles and open squares represent stratovolcanoes and the Asian continental margin due to the subduc- large-scale caldera volcanoes, respectively. tion of the Pacific plate and rifting and opening Tectonic shortening and coeval volcanism during the Quaternary 139 Table 1. Ages and magma volume (DRE: dense rock equivalent) for each eruptive episode. Strat., stratigraphical method; TL, thermo-luminescence dating; K-Ar, potassium argon dating; 14 C, radiocarbon dating; FT, fission track dating. (References o 4 () 6o0 a () 00 4 () 0 u (1) : f Quaternary Volcanoes in Japan 1999; (5) Nozawa 2001; (6) Aoki and Arai 2000; (7) Ohba et al. 2003; (8) Kano et al. 2007; (9) Yashima 1990; (10) Fujinawa et ai. 2001b; (11) Tsuchiya et al. 1997; (12) Mimura 2001; (13) Mimura and Kano 2000; (14） Umeda et al. 1999;(15) Fujinawa et al. 2001a; (16) Chiba and Kimura 2001;(17) Mimura 2002; (18) Yamamoto 2003, (19) Yamamoto and Komazawa 2004; (20) NEDO 1990). DRE Volcano Eruptive episode Latitude Longitude Age (Ma) Method Ref. (km²) Mutsuhiuchi-dake Older Mutsuhiuchi-dake 41.437 141.057 1.2-0.8 K-Ar (1) 5.9 Mutsuhiuchi-dake Younger Mutsuhiuchi-dake 41.437 141.057 0.8-0.5 K-Ar, FT (1) 3.6 Osorezan Kamabuse-yama 41.277 141.123 0.80-0.76 K-Ar, FT (2) 5.6 Osorezan Byobuyama-Asahinadake 41.277 141.123 0.7-0.5 K-Ar, FT (2) 3.2 Osorezan Pre-caldera pyroclastic fow 41.277 141.123 0.48-0.27 K-Ar, FT (2) 6.4 Osorezan Post-caldera pyroclastic fow 41.277 141.123 0.27-0.20 K-Ar, FT (2) 1.3 Osorezan Tsurugiyama 41.277 141.123 0.2-0.08 K-Ar, FT (2) 0.1 Hakkoda Hakkoda P.F. 1st. 40.667 140.897 0.76 Strat. (3) 17.8 Hakkoda South-Hakkoda 40.600 140.850 0.65-0.4 K-Ar (4) 52.4 Hakkoda Hakkoda P.F. 2nd. 0.4 K-Ar 40.667 140.897 (4) 17.3 Hakkoda North-Hakkoda 40.650 140.883 0.16-0 K-Ar (4) 30.4 Iwaki Iwaki 40.653 140.307 0.330-0.000 K-Ar (4) 37.2 Okiura AoniF.Aonigawa P.F. 40.573 140.763 ca. 1.7 K-Ar (5) 17.6 Okiura AoniF.Other P.F. 40.573 140.763 1.7-0.9 K-Ar (5) 3.7 Okiura Okogawasawalava 40.579 140.759 0.9-0.65 Strat. (5) 0.9 Okiura Okiuradacite 140.755 0.9-0.7 K-Ar 40.557 (5) 2.1 Ikarigaseki Nijikai Tuff 40.500 140.625 ca.2.0 K-Ar (4) 20.2 Ikarigaseki Ajarayama 40.490 K-Ar 140.600 1.91-1.89 (4) 2.1 Towada Herai-dake 40.450 141.000 5.1 Towada Ohanabe-yama 40.500 140.883 0.4-0.05 K-Ar (4) 8.9 Towada Hakka 40.417 140.867 1.4 14C Towada Towada Okuse 40.468 140.888 0.055 (4) 4.8 Towada Towada Ofudo 40.468 140.888 0.030 Strat. (6) 22.1 Towada Towada Hachinohe 40.468 140.888 0.015 Strat. (6) 26.9 Towada Post-calderacones 40.457 140.913 0.013-0 Strat. (4) 14.4 Tashiro Tashiro 40.425 140.413 0.600-0.470 K-Ar (4) 6.8 Tashiro Hirataki P.F. 40.420 140.413 0.020-0.020 Strat. (4) 0.3 Taira-Komagatake Taira-Komagatake 40.410 140.254 0.200-0.170 K-Ar (4) 2.3 Inaniwa Inaniwa 7.000-2.700 K-Ar 40.195 141.050 (4) 10.6 Nanashigure Nanashigure 40.068 K-Ar 141.112 1.06-0.72 (4) 55.5 Megata Megata 39.952 139.742 0.030-0.020 Strat. (4) 0.0 Toga Toga 39.950 139.718 ca. 0.42 FT/K-Ar (4) 1.2 Kampu Kampu 39.928 139.877 0.030-0.000 Strat. (4) 0.5 Moriyoshi Moriyoshi 39.973 140.547 1.07-0.78 K-Ar (4) 18.1 Bunamori Bunamori 39.967 140.717 1.2 K-Ar (4) 0.1 Akita- Yakeyama Akita-Yakeyama 39.963 140.763 0.5-0 K-Ar (4) 9.9 Nishimori/Maemori Nishimori/Maemori 39.973 140.962 0.5-0.3 K-Ar (4) 2.6 Hachimantai/Chausu Hachimantai 39.953 140.857 1.0-0.7 K-Ar (4) 5.5 Hachimantai/Chausu Chausu-dake 39.948 140.902 0.85-0.75 K-Ar (4) 13.7 Hachimantai/Chausu Fukenoyu 39.953 140.857 ca. 0.7 Strat. (4) 0.2 Hachimantai/Chausu Gentamri 39.956 140.878 (4) 0.2 Yasemori/Magarisaki-yama Magarisaki-yama 39.878 140.803 1.9-1.52 K-Ar (4) 0.3 Yasemori/Magarisaki-yama Yasemori 39.883 140.828 1.8 K-Ar (4) 0.9 Kensomori/Morobidake Kensomori 39.897 140.871 ca. 0.8 Strat. (4) 0.8 Kensomori/Morobidake Morobi-dake 39.919 140.862 1.0-0.8 Strat. (4) 2.5 Kensomori/Morobidake 1470 m Mt. lava 39.909 140.872 0.72 K-Ar (7) 0.1 Kensomori/Morobidake Mokko-dake 39.953 140.857 ca. 1.0 Strat. (4) 0.5 Tamagawa Welded Tuff Tamagawa Welded Tuffs R4 39.963 140.763 ca. 2.0 K-Ar (4) 83.2 Tamagawa Welded Tuff Tamagawa Welded Tuffs D 39.963 140.763 ca. 1.0 K-Ar (4) 32.0 140 Koji Umeda et al. Table 1. (Continued) DRE Volcano Eruptive episode Latitude Longitude Age (Ma) Method Ref. (km3) Nakakura/Shimokura Obuka-dake 39.878 140.883 0.8-0.7 K-Ar (4) 2.9 Nakakura/Shimokura Shimokura-yama 39.889 140.933 0.85-0.58 K-Ar (7) 0.4 Nakakura/Shimokura Nakakura-yama 39.888 140.910 0.85-0.58 K-Ar (7) 0.4 Matsukawa Matsukawa andesite 39.850 140.900 2.6-1.29 K-Ar (4) 11.6 Iwate/Amihari Iwate 39.847 141.004 0.2-0 K-Ar (4) 25.1 Iwate/Amihari Amihari 39.842 140.958 0.3-0.1 K-Ar (4) 10.6 Iwate/Amihari Omatsukura-yama 39.841 140.919 0.7-0.6 K-Ar (4) 3.3 Iwate/Amihari Kurikigahara 39.849 140.882 0.2 Iwate/Amihari Mitsuishi-yama 39.848 140.900 0.46 K-Ar (4) 0.6 Shizukuishi/Takakura Marumori 39.775 140.877 0.4-0.3 K-Ar (4) 2.4 Shizukuishi/Takakura Shizukuishi-Takakura-yama 39.783 140.893 0.5-0.4 Strat. (4) 5.2 Shizukuishi/Takakura Older Kotakakura-yama 39.800 140.900 1.4 K-Ar (4) 2.7 Shizukuishi/Takakura North Mikado-yama 39.800 140.875 0.3 Shizukuishi/Takakura Kotakakura-yama 39.797 140.907 0.6-0.5 K-Ar (4) 1.8 Shizukuishi/Takakura Mikado-yama 39.788 140.870 ca. 0.3 Strat. (4) 0.2 Shizukuishi/Takakura Tairagakura-yama 39.808 140.878 ca. 0.3 Strat. (4) 0.1 Nyuto/Zarumori Tashirotai K-Ar 39.812 140.827 0.3-0.2 (4) 0.6 Nyuto/Zarumori Sasamori-yama 39.770 140.820 0.23-0.1 K-Ar (4) 0.4 Nyuto/Zarumori Yunomori-yama 39.772 140.827 ca. 0.3 (4) 0.5 Nyuto/Zarumori Zarumori-yama 39.788 140.850 0.56 K-Ar (4) 0.9 Nyuto/Zarumori Nyutozan 39.802 140.843 0.58-0.5 K-Ar (4) 5.0 Nyuto/Zarumori Nyuto-kita 39.817 140.855 ca. 0.4 K-Ar (4) 0.1 Akita-Komagatake Akita-Komagatake 39.754 140.802 0.1-0 K-Ar (4) 2.9 Kayo Kayo 39.803 140.735 2.2-1.17 K-Ar (4) 5.9 Kayo KoJiromori 39.828 140.787 0.94 K-Ar (4) 0.3 Kayo Akita-Ojiromori 39.839 140.788 1.7 Strat. (4) 0.3 Daibutsu Daibutsu 39.817 140.517 2.34-2.16 K-Ar (4) 2.4 Tazawa Tazawa 39.723 140.667 1.80-1.40 FT (8) Innai/Takahachi Takahachi-yama 39.755 140.655 1.7 K-Ar (4) 0.0 Innai/Takahachi Innai 39.692 140.638 2.0-1.6 K-Ar (4) 0.5 Kuzumaru Aonokimori andesites 39.543 140.983 2.06 K-Ar (9) 0.3 Yakeishi Yakeishidake 39.161 140.832 0.7-0.6 K-Ar (4) 9.5 Yakeishi Komagatake 39.193 140.924 ca. 1.0 K-Ar (4) 7.6 Yakeishi Kyozukayama 39.178 140.892 0.6-0.4 K-Ar (4) 5.7 Yakeishi Usagimoriyama 39.239 140.924 0.07-0.04 K-Ar (4) 2.3 Chokai Shinsan Lava fow 39.097 140.053 0.02-0 Strat. (4) 0.6 Chokai Higashi Chokai 39.097 140.053 0.02-0.02 K-Ar (4) 3.3 Chokai Nishi Chokai 39.097 140.020 0.09-0.02 Stra. (4) 0.6 Chokai Nishi Chokai II 39.097 140.020 0.13-0.01 K-Ar (4) 16.0 Chokai Old Chokai 39.103 140.030 0.55-0.16 K-Ar (4) 50.9 Chokai Uguisugawa Basalt 39.103 140.030 0.6-0.55 K-Ar (4) 0.8 Chokai Tengumari volcanics 39.103 140.031 0.6-0.55 K-Ar (4) 8.4 Kobinai Kobinai 39.018 140.523 1.0-0.57 K-Ar, FT (4) 2.3 Takamatsu/Kabutoyama Kabutoyama Welded Tuffs 1.16 39.025 140.618 TL (4) 3.2 Takamatsu/Kabutoyama Kiji-yama Welded Tuffs 39.025 140.618 0.30 K-Ar (4) 5.1 Takamatsu Takamatsu 38.965 140.610 0.3-0.27 K-Ar (4) 3.8 Takamatsu Futsutsuki-dake 38.961 140.661 ca. 0.3 (4) 0.8 Kurikoma Tsurugi-dake 38.963 140.792 0.1-0 K-Ar (10) 0.2 Kurikoma Magusa-dake 38.968 140.751 0.32-0.1 K-Ar (10) 1.5 Kurikoma Kurikoma 38.963 140.792 0.4-0.1 K-Ar (10) 0.9 Kurikoma South volcanoes 38.852 140.875 ca. 0.5 K-Ar (10) 0.3 Kurikoma Older Higashi Kurikoma 38.934 140.779 ca. 0.5 K-Ar (10) 2.2 Kurikoma Younger Higashi Kurikoma 38.934 140.779 0.4-0.1 K-Ar (10) 0.7 Tectonic shortening and coeval volcanism during the Quaternary 141 Table 1. (Continued) DRE Volcano Eruptive episode Latitude Longitude Age (Ma) Method Ref. (km²) Mukaimachi Mukaimachi 38.770 140.520 ca. 0.8 K-Ar (4) 12.0 Onikobe Shimoyamasato Tuff 38.830 140.695 0.21 FT (4) 1.0 Onikobe Onikobe Central cones 38.805 140.727 ca.0.2 TL (4) 1.1 Onikobe Ikezuki Tuff 38.830 140.695 0.3-0.2 FT (4) 17.3 14C Naruko Naruko Central cones 38.730 140.727 0.045 (4) 0.1 Naruko Yanagizawa Tuff 38.730 140.727 0.045 FT (4) 4.8 Naruko Nizaka Tuff 38.730 140.727 0.073 FT (4) 4.8 Hijiori Hijiori Pyroclastic fow 38.610 140.159 ca. 0.01 Strat. (4) 0.5 Hijiori Komatsubuchi lava dome 38.613 140.171 ca. 0.01 Strat. (4) 0.0 Gassan Ubagatake 38.533 140.005 0.400-0.300 K-Ar (4) 2.7 Gassan Yudonosanlavas/pyroclastics 38.534 139.988 0.800-0.700 K-Ar (4) 5.7 Gassan Gassan 38.550 140.020 0.500-0.400 K-Ar (4) 13.7 Funagata Izumigatake 38.408 140.712 1.45-1.14 K-Ar (4) 2.3 Funagata Funagatayama 38.453 140.623 0.85-0.56 K-Ar (4) 19.0 Yakuraisan Yakuraisan 38.563 140.717 1.65-1.04 K-Ar (11) 0.2 Nanatsumori Nanatsumori lava K-Ar 38.430 140.835 2.3-2.0 (12) 0.5 Nanatsumori Miyatoko Tuffs 38.428 140.793 ca. 2.5 Strat. (4) 6.1 Nanatsumori Akakuzure-yama lava 38.433 140.768 1.6-1.5 Strat. (4) 1.5 Kamikadajin lava Nanatsumori 38.447 140.772 1.6-1.5 K-Ar (4) 0.8 Shirataka Shirataka 38.220 140.177 K-Ar 1.0-0.8 (13) 3.8 Adachi Adachi 38.218 140.662 0.08 FT (4) 0.9 Gantosan Gantosan 38.195 140.480 0.4-0.3 K-Ar (4) 4.6 Kamuro-dake Kamuro-dake 38.253 140.488 1.67 K-Ar (12) 5.7 Daito-dake ca. 1.0 Strat. Daito-dake 38.316 140.527 (4) 5.7 Sankichi-Hayama Sankichi-Hayama 38.137 140.315 2.400-2.300 K-Ar (4) 2.2 Ryuzan Ryuzan 38.181 140.397 1.1-0.9 K-Ar (4) 4.6 Zao Central Zao lst. 38.133 140.453 1.46-0.79 K-Ar (14) 0.8 Zao Central Zao 2nd. 38.133 140.453 0.32-0.12 K-Ar (14) 15.2 Zao Central Zao 3rd. 38.133 140.453 0.03-0 K-Ar (14) 0.0 Zao Sugigamine 38.103 140.462 1.0 K-Ar (14) 9.9 Zao Fubosan/byobudake 38.093 140.478 0.31-0.17 K-Ar (14) 15.2 Aoso-yama Gairinzan 38.082 140.610 0.7-0.4 K-Ar (12) 6.1 Aoso-yama Central Cone 38.082 140.610 0.40-0.3 K-Ar (4) 3.0 Azuma Azuma Kiteilava 37.733 140.247 1.3-1.0 K-Ar (4) 24.7 Azuma Higashi Azumasan 37.710 140.233 0.7-0 K-Ar (4) 22.8 Azuma Nishi Azumasan 37.730 140.150 0.6-0.4 K-Ar (4) 7.2 Azuma Naka Azumasan 37.713 140.188 0.4-0.3 K-Ar (4) 4.6 Nishikarasugawaandesite Nishikarasugawa andesite 37.650 140.283 ca. 1.5 K-Ar (4) 1.9 Adatara Adatara Stage 1 37.625 140.280 0.55-0.44 K-Ar (15) 0.3 Adatara Adatara Stage 2 37.625 140.280 ca. 0.35 K-Ar (15) 0.4 Adatara Adatara Stage 3a 37.625 140.280 ca. 0.20 K-Ar (15) 2.0 Adatara Adatara Stage 3b 37.625 140.280 0.12-0.002 K-Ar (15) 0.3 Sasamori-yama Sasamari-yama andesite 37.655 140.391 2.5-2 K-Ar (4) 0.4 Bandai Pre-Bandai 37.598 140.075 ca. 0.7 K-Ar (4) 0.1 Bandai Bandai 37.598 140.075 0.3-0 Strat. (16) 14.0 Nekoma Old Nekoma 37.608 140.030 1.0-0.7 K-Ar (17) 11.4 Nekoma New Nekoma 37.608 140.030 0.5-0.4 K-Ar (17) 0.9 Numazawa Shirifukitoge P. 37.452 139.577 0.11 FT (18) 0.7 Numazawa Mukuresawa L. 37.452 139.577 0.071 TL (18) 0.1 14C Numazawa Mizunuma P. 37.452 139.577 0.045 (18) 1.0 Numazawa Sozan L. 37.452 139.577 0.043 FT (18) 0.3 14C Numazawa Numagozen P. 37.452 139.577 0.0198 (18) 0.0 14C Numazawa Maeyama L. 37.452 139.577 0.02 (18) 0.3 142 KojiUmeda et al. Table 1.(Continued) DRE Volcano Eruptive episode Latitude Longitude Age (Ma) Method Ref. (km3) Numazawa Numazawako P. 37.452 139.577 0.005 14C (18) 2.0 Sunagohara Sunagohara 37.457 139.684 0.290.22 FT (19) 2.7 Sumon Sumon 139.140 2.401.75 KAr (4) 22.0 Asakusa Asakusa 37.340 139.237 1.641.54 KAr (4) 4.6 Kasshi/Oshiromori Kasshi 37.184 139.973 0.1 Kasshi/Oshiromori Oshiromori 37.199 139.970 0.7 Kasshi/Oshiromori Matami-yama 37.292 139.886 0.94 KAr (20) 80 Kasshi/Oshiromori Naka-yama 37.282 139.899 0.0 Shirakawa Kumado P.F. 37.242 140.032 1.31 KAr (4) 19.2 Shirakawa Tokaichi A.F. tuffs 37.242 140.032 1.311.24 Strat. (4) 12.0 Shirakawa Ashino P.F. 37.242 140.032 1.2 FT (4) 19.2 Shirakawa Nn3 P.F. 37.242 140.032 1.201.17 Strat. (4) 0.0 Shirakawa Kinshoji A.F. tuffs 37.242 140.032 1.201.18 Strat. (4) 9.0 Shirakawa Nishigo P.F. 37.252 139.869 1.11 FT (4) 28.8 Shirakawa Tenei P.F. 37.242 140.032 1.06 Strat. (4) 7.7 Nasu Futamata-yama 37.244 139.971 0.14 KAr (4) 3.2 Nasu Kasshiasahi-dake 37.177 139.963 0.60.4 KAr (4) 12.3 Nasu Sanbonyari-dake 37.147 139.965 0.40.25 KAr (4) 5.5 Nasu Minamigassan 37.123 139.967 0.20.05 KAr (4) 8.7 Nasu Asahi-dake 37.134 139.971 0.20.05 KAr (4) 4.6 Nasu Chausu-dake 37.122 139.966 0.040 KAr (4) 0.3 of the Japan Sea.The main rifting started at contributedtothegentleupliftofthecentral ~23 Ma (Taira 2001).Between 21 and 18 Ma, mountains range (Yoshida 2001). rifting was accompanied by significant counter- Tectonic shortening of the crust due to E-W clockwise rotation of the NE Japan arc (Jolivet compressionbecame apparentataboutthe et al. 1994). Owing to the cessation of the open- Pliocene to Quaternary boundary,which may be ing of the Japan Sea,the extensional stress field associatedwith theincreased motion of thePacific changed at about 13 Ma.From the Middle Miocene Plate between 5 and 2 Ma (Cox and Engebretson to the Pliocene,the tectonics is characterized by 1985;Pollitz 1986).A compressional stress field very weak crustal deformation under the moderate during the Quaternary is responsible for the devel- regional stress field related to the convergence of opment of two narrow uplift zones oriented in a the Pacific plate (Sato 1994). The maximum hori- NS direction, in the NE Japan arc: the Ou Back- zontal stress oriented in the NE or ENE direction bone Range (fore-arc) and the Dewa Hills (rear- was manifested during this period. This is one of arc). They appear to be an active pop-up structure the reasons why the SW migration of the Kuril bounded by opposite-facing reverse faults accom- sliver duetotheobliqueconvergencealongthe modating <5 mm/yr of EW shortening across Kuril arc results from a NE or ENE trending max- the range (Hasegawa et al. 2005). Based on the imum compression (e.g., Otsuki 1990). In relation subsurface geology and deformation of river ter- to the stress change from extensional to a slightly races,the initiation time of reverse faulting was compressive stress, a major variation in the volcan- estimated at several sites in the NE Japan. These ism in the NE Japan arc changed from rift-type results suggest that reverse faulting started on the into island-arc type in the last 13 Ma (Sato 1994; rear-arc side between 3.4 and 2.4 Ma (Awata and Yoshida 2001), resulting in changes in the loca- Kakimi 1985), and in the fore-arc side between 0.9 tion and orientation of the volcanic front (Ohguchi and 0.5 Ma (Otsuki et al.1977),corresponding to et al.1989).Submarine basaltic to rhyolitic vol- the onset time of uplift of the Dewa Hills and the canism was predominant from 13 to 8 Ma. Many Ou Backbone Range.The compressionalregime Valles-type caldera (volcanic crater formed by col- reactivated originally normalfaults related to the lapse during an ignimbrite eruption） formed on the extensional back-arc rifting as reverse faults and centralmountains range along the arc since ca. these accommodate much of the ongoing shorten- 8Ma.The dome-like structures around the calderas ing across the arc.Andesite stratovolcanoes started indicate that felsic magmatism in the shallow crust toform between 2.5 and 1.5 Ma,but these are Tectonic shorteningand coevalvolcanism during theQuaternary 143 relatively rare compared with those younger 1.5 Ma are: stratovolcanoes, 1.9 g/cm²;non-welded pyro- volcanoes (Kimura and Yoshida 2006).Neverthe- clastic flow deposits,1.2g/cm²;welded pyroclastic less,it remains obscure as to when the calc-alkaline fow deposits,1.6 g/cm²;pyroclastic fall deposits, andesite volcanism, derived from the mixing between 1.5 g/cm² (Nakamura 1964; Smith 1979;Yoshida the felsic and mafic magmas (e.g., Sakuyama and and Takahashi 1991). Nesbitt 1986),has been established under Quater- Table 1 presents the locations, active time spans nary orthogonal convergence in the NE Japan arc. (range),and magma volumes of the individual eruptive episodes during the Quaternary,includ- ing the Gelasian Stage.Sixty-nine volcanoes com- 3.Estimation of timingandvolume posed of 176 eruptive episodes can be identified foreacheruptiveepisode along the volcanic arcbetween Mutsuhiuchi-dake volcano and Nasu volcano (figure 1). Although 34 Quaternary volcanoes havebeenrec- ognized along the volcanic arc between Mutsuhiuchi- dake volcano and Nasu volcano in the NE Japan 4.Time-volumerelationshipsforNEJapan arc (Ono et al.1981),they are predominantly volcanic arc since 2.0Ma polygenic, with new edifices forming at new loca- tions and evolving into polygenic volcano centres. In order to elucidate temporal variations in the long- Before refining the sequence of volcanism during term magma discharge rate for all of the NE Japan the Quaternary,we subdivided individual volca- noes identified by Ono et al. (1981) into as small a units as possible. The term ‘eruptive episode’ used 1000 herein refers to any unit that formed due to a ser- a ies of eruptions from the same conduit or edifice,in- 800 cluding monogenic volcanoes (volcano constructed during a single phase of eruptive activities) that 600 formed in a relatively short time period and poly- genetic volcanoes (resulting in from more than one 400 formation process) that developed over several tens to hundreds of thousands of years. The active time 200 span of eruptive episodewithseveralradiomet- ric ages can be estimated by their median age. 2.0 1.5 1000 1.0 0.5 Radiometric ages that are inconsistent with the stratigraphic data were discarded. b The magma volume discharged for each eruptive 800 episode was calculated as follows: Assuming that 600 a edifice isrepresented by a cone with a radius of basal circle (R)and height between the base and 400 apex (H),volume (V) is calculated by the follow- ing equation: V =(1/3) HπR².When it is repre- 200 sented by an elliptical cone, the radius of a basal circle (R') with the same area as the basal ellipse is found, and then the R'is applied in the above equa- 1200 2.0 1.5 1.0 0.5 tion (Umeda et al.1999).The relief of the underly- 1000 C ing basement can be considered.For example,when the basement of the conical edifice is uplifted in the 800 centralpart to form a raised bottom,the excessive volume is subtracted from the conical volume. The 600 volume of lava flows or pyroclastic flows is calcu- 400 lated by multiplying their basal area and average thickness.When it is difficult to estimate the aver- 200 age thickness,their distribution area is subdivided into smaller areas by taking into account basement 2.0 1.5 1.0 0.5 relief, and the products of their area and average Age (Ma) thickness in the respective subdivided areas are totalled. The DRE of erupted volumes was calcu- Figure 2. Timevolume relationships for the NE Japan vol- lated by the productofvolume anddensity of the canic arc since 2.0 Ma. (a) Volcanism associated with stra- tovolcanoes,(b)felsic volcanism associated with the large respective volcanic products.Density values used calderas, and (c) total erupted magma volume. 144 Koji Umeda et al. arc, the magma volume erupted every 100,000-year There are two fundamental types of time volume period (long-term discharge rate of magma) was relationships for volcanic fields: volume-predictable calculated for each volcano since 2.0 Ma. In the behaviour (linear relationship between the cumu- case of South-Hakkoda eruptive episode between lative volume of an eruptive episode plus prior 0.65 and 0.40 Ma belonging to Hakkoda volcano episodes and its timing) and time-predictable (table 1), the erupted volume was estimated to behaviour (linear relationship between the tim- be 52.4 km? magma from which 10.4, 21.0 and ing of an episode and the cumulative volume of d pn s prior episodes only). A plot of total volume-age of 0.7-0.6, 0.6-0.5 and 0.5-0.4 Ma, respectively. relationship throughout the NE Japan arc sug- Figure 2 shows plots of cumulative volume asso- gests that it appears to change somewhat from the ciated with large, felsic calderas and andesitic time-predictable trend into the volume-predictable stratovolcanoes as a function of their ages. It indi- trend around 1.0 Ma(figure 2). Valentine and cates that eruptive volume fux of andesitic stra- Perry (2007） suggested that eruption marked by tovolcanoes increased significantly after 1.1 Ma. the volume-predictable trend occurs in magmati- Although fux of eruptive volume from large, fel- cally controlled fields (high magma fuxes), where sic calderas increase slightly between 1.4 and pressures build up in magma reservoirs, driven 1.0 Ma, the slope, except the slightly anomalous primarily by process such as melt accumulation, period indicates a constant fux of ~0.1 km?/ky. fractionation, and concentration of volatiles. Accordingly, volcanism can be divided into two Most of the volcanoes exist on uplifted regions stages: Stage 1 (before ca. 1.0 Ma) is charac- forming a topographic high known as the Ou Back- terized by the predominance of large-volume fel- bone Range. In addition, E-W volcanic zones have sic eruptions forming large calderas, Stage 2 (ca. been developed with several volcanoes in the rear 1.0 Ma onwards), is dominated by calc-alkaline arc (figure 1). It should be noted that the distri- andesite stratovolcano-building eruption, except bution of volcanoes formed in Stage 1 is limited on for Towada volcano. Stage 2 is marked by a sig- the Ou Backbone Range. However, Stage 2 is char- nificant increased magmatism in the entire NE acterized by volcanism extended to the rear arc, Japan arc. Felsic volcanism, dominant in Stage 1, suggesting that Quaternary NE Japan volcanism, is believed to have occurred under neutral to weak stratovolcanoes arranging in E-W direction per- compressional stress fields associated with gentle pendicular to the volcanic arc, have been estab- uplifting of volcanic arc (Yoshida 2001). lished since ca. 1.0 Ma (figure 3). 42 41 1 Ot 39- 39 2550 37 37 138 139 140 141 142 143138 139 140 141 142 143 Stage 1 (2.0-1.0 Ma) Stage 2 (1.0 Ma-present) n ( ( ) active faults. Tectonic shortening and coeval volcanism during the Quaternary 145 5. Discussion and conclusions fault development, which had straight traces at the free surface. When shortening of the silica powder As mentioned in the previous section, the NE and injection of the oil were simultaneous, magma Japan arc has been undergoing E-W compression formed a basal sill and then typically rose along without any significant oblique component during a thrust fault. These results imply that, theoreti- the Quaternary, and reverse faulting has occurred cally, compression favours the emplacement of sills across the entire volcanic arc since about 1.0 Ma, (Hubbert and Willis 1957), but thrust faults then resulting in the predominance of the N-S trending provide vertical tension fractures serving as con- thrusts with crustal shortening. Contemporane- duits for magma in the shallow portion of the hang- ous changes in volcanism occurred as well; around ing wall, demonstrating that magma can propagate 1.0 Ma, andesite volcanism with stratovolcano- to the Earth's surface in a tectonic regime char- forming eruptions was initiated. Moreover, magma acterized by contracting deformation. These struc- discharge rate in the entire volcanic arc has rapidly tures can be attributed to superposition of a local increased since ca. 1.o Ma. At the same time, the load, due to the uplifted zone, on the regional stress magma underwent a systematic change in chemi- feld (Molnar 1986). Therefore, apparent thrust cal composition. A significant volume of medium-K faulting throughout the Ou Backbone Range since andesite has been erupted along the Ou Backbone ca. 1.0 Ma could facilitate the effective transfer of Range since 1.0 Ma to 0.7 Ma, together with sub- magma from the crust and localize newly formed ordinate low-K andesite (Ban et al. 1992). What volcanoes close to thrust faults on Ou Backbone are the processes linked to volcanic and tectonic Range. changes around 1.0 Ma in the NE Japan arc? Consequently, volcanism associated with andes- Several researchers insist that the Pacific plate ite stratovolcanoes, and thrust faulting due to tec- underwent significant changes in the amount and tonic shortening has been clearly active since ca. absolute direction of motion between 5 and 2 Ma 1.0 Ma, in the NE Japan arc. This is because the (Cox and Engebretson 1985; Pollitz 1986). These increased melt production rate within the mantle studies agree on a significant increase in the sub- wedge and compressional stress regime is ascribed duction rate of the Pacific plate, estimated, in to a significant increase in the subduction rate of the NE Pacific, between ~30 and ~70%. In con- the Pacific plate between 5 and 2 Ma. The contrac- sequence, the NE Japan arc has been undergo- tional tectonic settings that form a basal sill for ing active compressional inversion with shortening magma prevent magma from reaching the Earth's intensifying over 2.0 Ma. surface. Nevertheless, some volume of magma from However, despite compressional stress regime the basal sill could be extruded onto the sur- (o = 03 <o2 <o1),a large amount of magma face where uplifted mountains, bounded by reverse has been extruded since ca. 1.0 Ma. One of the faults, perpendicular to the volcanic arc, have been plausible explanations for this contradiction may established since ca. 1.0 Ma. Rather slow ascent be the increase in magma generation in the man- tle wedge. Numerical simulations considering fuid entiated felsic magma within the upper crust to migration and melting in the mantle wedge above mix with primary magma arising from the deeper a subducting plate indicate that melt production potion, such as lower crust and/or mantle wedge, rates increase with increasing convergence rate resulting in the production of calc-alkaline andesite (Cagnioncle et al. 2007). In the Cascade volcanic volcanism. arc, the convergence rate of the Juan de Fuca plate to the North American plate is thought to Acknowledgements control the change in eruption rate (Priest 1990). Therefore, despite the overall compressive setting. The authors would like to thank T Ohba,1 M along the NE Japan arc, the increase in volume of Sasaki and K Akaishi for helpful discussions. Care- magma could be due to the increase in degree of ful reviewing of R I Tilling, S J Day, G F McCrank partial melting in the mantle wedge promoted by and the anonymous reviewers has greatly improved significantly faster subducting Pacific plate. the paper. In addition, recent experimental modelling pro- vided insight into the upwelling process of magma References within brittle crust under a compressive stress regime (Galland et al. 2007). These researchers Acocella V, Yoshida T, Yamada R and Funiciello F 2008 employed (1) silica powder to replicate a brit- Structural control on Late Miocene to Quaternary volcan- tle crust and (2) liquid vegetable oil representing ism in the NE Honshu arc, Japan; Tectonics 27 TC5008, doi: 10.1029/2008TC002296. magma, to model materials. 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Umenda (2013) Tectonic shortening and coeval volcanism during.txt
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J. metamorphic Geol., 1998, 16,129–140
Geotectonic subdivision and areal extent of the Sangun belt, Inner
Zone of Southwest Japan
Y. NISHIMURA
Department of Earth Sciences, Faculty of Science, Yamaguchi University, Yamaguchi 753, Japan
ABSTRACT The Sangun belt has long been considered to be a major coherent glaucophanitic terrane of Permian to
Triassic age, and to be paired with the low- P/THida belt to the north. However, recent progress in
geochronology, metamorphic geology, and tectonics has revealed that the belt is in fact comprised of twogeologic units of di Verent ages and with contrasting conditions of formation. The older unit is named
the Renge belt and the younger the Suo belt. The Renge belt is the oldest of the high- P/Tmetamorphic
belts in the Japanese Islands and extends from northern Kyushu, through the San-in coastal regions, tothe Hida marginal belt. It is characterized by 330–280 Ma ages and the association of glaucophane–schist
to epidote–amphibolite facies schists. The Renge belt is also typically associated with meta–ophiolite
sequences (470–340 Ma) including serpentinite. The Suo belt is characterized by 230–160 Ma high- P/T
schists closely related to weakly metamorphosed Permian accretionary rocks of the Akiyoshi belt.
Metamorphic facies series is from the prehnite–pumpellyite facies through the pumpellyite–actinolite and
glaucophane–schist facies to the epidote–amphibolite facies. The belt is widespread in west Kinki to northand central Kyushu via Chugoku, but also stretches further to the southwest and is present in the Ishigaki-
Iriomote Islands of the southern Ryukyu Arc. Throughout this belt, there are scattered small blocks or
lenses of meta–ophiolite, whose K–Ar ages of relict hornblendes are 590 to 220 Ma. Bounded by low-angle faults and thrusts, both belts define subhorizontal nappes dipping gently north. The geotectonic
framework in the Inner Zone of Southwest Japan is made up of, from north to south, the Hida-Oki,
Renge, Akiyoshi, Suo, Maizuru plus ultra-Tamba, Mino-Tamba, and Ryoke belts, with a tectonicallydownward-younging polarity. This has resulted from stepwise accretions during Palaeozoic to
Mesozoic time.
Key words: high-P/Tmetamorphism; radiometric ages; Renge belt; Sangun belt; Suo belt.
heterochronous belts formed during separate metamor-INTRODUCTIONphic events (Nishimura et al., 1983b: Shibata &
Nishimura, 1989). Nishimura (1990) proposed a three- The Sangun metamorphic belt proposed by Kobayashi
(1941) has been considered to be a major coherent fold division of the belt, and assigned new names to
their divisions: the Sangun-Renge belt ( c. 300 Ma), the metamorphic terrane of Permian to Triassic age, and
to constitute the axial core of the Akiyoshi orogen Suo belt ( c. 220 Ma), and the Chizu belt ( c. 180 Ma),
from oldest to youngest. The concept of a regional along with the Hida gneiss in the Inner Zone of
Southwest Japan. Miyashiro (1961) interpreted the Sangun-Renge belt has been generally accepted.
However, the distinction between the Suo and Chizu Sangun belt as a high- P/Tmetamorphic belt that is
paired with the low- P/THida metamorphic belt to belts is less widely accepted, owing to the lack of
geotectonic evidence and the narrow age interval the north. These views have long been accepted among
Japanese geologists. Since the late 1970s, microfossil between the two (e.g. Isozaki & Maruyama, 1991;
Banno & Nakajima, 1992; Hayasaka et al., 1995). biostratigraphy and tectonic studies have achieved a
major breakthrough in the understanding of Japanese There is, therefore, a need to re-examine the nature of
the Suo and Chizu belts. basement geology. The framework of interpretation
has changed from geosynclinal orogeny to plate In this paper, the author aims to clarify the
metamorphic geology of the Sangun belt, and to tectonics or accretion tectonics. As a result, the Sangun
belt has been recognized as representing a subduction- establish the geotectonic subdivision and areal extent
of the belt, based on new data obtained by him and related high- P/Tmetamorphic belt that formed along
the early Mesozoic convergent margin of east Asia his collaborators. It has been proposed that the
Kurosegawa belt in the Outer Zone of Southwest (Uyeda & Miyashiro, 1974; Maruyama & Seno, 1986).
However, following recent systematic studies of Japan represents a tectonic outlier of the pre-Jurassic
rocks of the Inner Zone (Isozaki & Itaya, 1991; Isozaki geochronology and metamorphic geology, it has
become clear that the Sangun belt is not a major et al., 1992). However, this problem will not be dealt
with here. In order to avoid confusion, this paper uses coherent high- P/Tmetamorphic belt, but consists of
129 © Blackwell Science Inc., 0263-4929 /98/$14.00
Journal of Metamorphic Geology , Volume 16, Number 1, 1998, 129–140
130 Y. NISHIMURA
the term ‘Sangun’ in a wide sense, and the term ‘belt’ pelitic and basic schists accompanied by psammitic,
siliceous and calcareous schists. Meta–ophiolitic rocks, designates an elongated metamorphic area. All the
ages cited are calculated by use of the decay constants such as metagabbro, amphibolite and serpentinite, also
occur in subordinate amounts. These rocks are inti- recommended by the IUGS Subcommission on
Geochronology (Steiger & Ja ¨ger, 1977) and the mately associated with non- to weakly metamorphosed
Palaeozoic formations and the Yakuno ophiolitic geologic time scale used is that of Harland et al. (1990).
complex. Abundant Cretaceous granitic rocks intruded
these rocks and caused contact metamorphism over aMETAMORPHIC GEOLOGY ANDbroad region.GEOCHRONOLOGYMetamorphic zonal mapping was carried out in
several areas, metamorphic rocks of which were Figure 1 shows the distribution of the Sangun meta-
morphic rocks in the Inner Zone of Southwest Japan classified into the following four zones (Hashimoto,
1968, 1972; Nishimura, 1971; Nishimura et al., 1977). along with radiometric ages obtained by Shibata &
Nishimura (1989). The rocks occur in many areas With rising temperature, the metamorphic rocks
change from the pumpellyite–chlorite zone through scattered from north and central Kyushu to west Kinki
via Chugoku. Although individual areas are small and pumpellyite–actinolite and epidote–glaucophane zones
into the barroisite zone. The lower temperature zones, isolated, they are zonally arranged in an ENE-WSW
direction. Geologists have therefore surmised the pumpellyite–chlorite to lower pumpellyite–actinolite,
correspond to the Palaeozoic formations and the existence of a major belt of regional metamorphism.
The Sangun metamorphic rocks consist mostly of Yakuno ophiolitic complex, while the other three zones
Fig. 1. Geological map of the Sangun belt in the Inner Zone of Southwest Japan showing three groups of radiometric ages from
the belt (after Shibata & Nishimura, 1989) and related areas (see text). Geochronological data and symbols: numeral =radiometric
age in Ma, [] =data from same specimen, K =K–Ar method, R =Rb–Sr method, h =barroisite, m =phengitic mica, w =whole-rock
isochron, upper row =Sangun-Renge belt (Renge belt in Figs 8 & 9), middle row =Chizu belt (Suo belt in Figs 8 & 9), lower
row=Suo belt (Suo belt in Figs 8 & 9). Geological abbreviations: Sib =San-in branch of the Sangun belt, Syb =San-yo branch of
the Sangun belt, N-zone =north zone (Akiyoshi belt in Fig. 9), M-zone =middle zone (Maizuru plus ultra-Tamba belts in Fig. 9),
Ry=Ryoke metamorphic belt, Tm =Tamba zone (Mino-Tamba belt in Fig. 9), Nz =Nagato tectonic zone.
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GEOTECTONIC SUBDIVISION OF THE SANGUN BELT 131
are the Sangun metamorphic rocks in a wide sense. metamorphic rocks as a single metamorphic unit
(Kizaki & Takayasu, 1975). Recently Isozaki & Based on the above facies series, it is thought that the
rocks belong to a single cycle of glaucophanitic Nishimura (1989) and Ujiie & Nishimura (1992) have
revised this as follows. metamorphism.
Shibata & Nishimura (1989) systematically deter- The Fusaki Formation is dominated by a chaotic
complex of various allochthonous blocks in a muddy mined 44 radiometric ages from 18 areas in the Sangun
belt, as shown in Fig. 1. The majority of the ages are matrix. The blocks consist of chert, siliceous mudstone,
sandstone and limestone. Conodonts, radiolarians and those of phengitic micas in pelitic and psammitic
schists utilizing K–Ar and Rb–Sr methods. Rb–Sr smaller foraminifera found in the allochthonous blocks
indicate latest Carboniferous to Early Jurassic ages whole-rock isochron ages were also calculated. In
determining radiometric ages, as many individual (Isozaki & Nishimura, 1989). Although the age of the
muddy matrix has not yet been determined, it is specimens as possible were measured by both K–Ar
and Rb–Sr methods, and the discrepancies resulting reasonable to suppose that the Fusaki Formation was
formed as a part of an accretionary complex in a from the two di Verent methods were carefully exam-
ined. The following age data in Fig. 1 are rejuvenated Middle Jurassic subduction zone. Following mixing in
the subduction zone, the formation was metamor- ages due to contact metamorphism confirmed petro-
graphically: 26 Km, 119 Rm, 129 Km and 158 Rm in phosed to at least slate to phyllite grade (zone I in
Fig. 2). The ages of the metamorphism as determined Wakasa (Hattou Fm.), 259 Km in Wakamiya, and 163
Km in Kurume. Those K–Ar ages that are younger by K–Ar ages of recrystallized phengitic micas in seven
pelitic rocks (Fig. 2; Ujiie & Nishimura, 1992) fall in than Rb–Sr ages by more than 10 My for the same
specimen given in Fig. 1 are also interpreted to be the range of 145–130 Ma (Early Cretaceous).
The Tomuru Formation is composed mainly of basic rejuvenated ages (Shibata & Nishimura, 1989). They
are as follows: 154 Km in Wakasa (Hattou Fm.), 120 and pelitic schists, accompanied by siliceous and
psammitic schists. Small blocks of meta–ophiolite, such Km in Asahi-cho, 165 Km in Tsukita, 206 Km in
Yamaguchi, 272 Km in Wakamiya, 184 Km in Kurume, as metagabbro and serpentinite, also occur concordantly
in the schists of zone II. These rocks have been and 193 Km in Yamaga.
After excluding these rejuvenated ages, the age data completely recrystallized and are divided into three
zones in accordance with the mineral parageneses in in Fig. 1 range from 308 to 174 Ma, i.e. from the
Carboniferous to Jurassic. The available data in Fig. 1 the basic schists (Fig. 2; Nishimura et al., 1983a). Zone II
is characterized by the pumpellyite–glaucophane assem- were tentatively divided into three groups, each of
which is distributed in a certain limited region blage, lawsonite and aragonite also being stable
(Ishizuka & Imaizumi, 1988). Metagabbro is metamor- (Nishimura, 1990). They are the Sangun-Renge belt
(310–280 Ma: c. 300 Ma), the Suo belt (230–210 Ma: phosed to the same mineral assemblages as the basic
schists. Zone III is marked by the disappearance of c. 220 Ma) and the Chizu belt (200–170 Ma: c.
180 Ma). pumpellyite, lawsonite and aragonite, and is charac-
terized by the epidote–glaucophane mineral assemblage.Zone IV is characterized by the appearance of barroisiteRECENT PROGRESS: AREAL EXTENTin basic schists and of garnet in pelitic schists. The
boundaries between zones II, III and IV are nearly Because the three-fold division of the Sangun belt was
proposed by Nishimura (1990), much more data are parallel to the bedding schistosity with no apparent
tectonic gap. The metamorphic grade increases from available concerning metamorphic petrology, K–Ar
dates, and tectonics of the belt and its related areas. the lower to the upper stratigraphic levels in the
Tomuru Formation, representing an inverted metamor- The new results require a reconsideration of the
previously proposed three-fold division of the Sangun phic gradient. Because these rocks have been completely
recrystallized under a progressive metamorphism of belt.
high-P/Ttype, their depositional ages remain unknown.
Recently K–Ar phengitic mica ages were determined forTomuru metamorphic rocks15 pelitic schists. They cluster around 220–190 Ma
(Fig. 2; Ujiie & Nishimura, 1992), which is a range Pre-Tertiary rocks occur widely from Ishigaki to
Iriomote Islands in the southern Ryukyu Arc, 1000 km more restricted than the previously reported ones (240–
160 Ma; Shibata et al., 1968, 1972; Nishimura et al., southwest of Kyushu (Figs 2 & 3). The rocks can be
divided into two formations based on their geologic 1983a; Faure et al., 1988b). Four K–Ar ages reported
by Nishimura et al. (1983a) are younger by 40 to 50 My setting and lithofacies. High- P/Tmetamorphic rocks
are designated as the Tomuru Formation, whereas than new K–Ar ages, even although the samples from
the same exposure were analysed, and must hence be a weakly metamorphosed olistostromal complex is
referred to as the Fusaki Formation (Foster, 1965). eliminated from consideration.
To recapitulate, the Fusaki Formation di Vers from The Tomuru Formation is thrust over the Fusaki
Formation. Both formations were previously combined the Tomuru Formation in its original depositional
environment and metamorphic history. The younger as the Ishigaki Group (Foster, 1965) or the Yaeyama
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132 Y. NISHIMURA
Fig. 2. Regional distribution of metamorphic zones and K–Ar ages from pre-Tertiary rocks in the Ishigaki-Iriomote Islands (after
Ujiie & Nishimura, 1992). The location of the islands is shown in Fig. 3. Star represents Early Jurassic radiolarians, diamondTriassic radiolarians and conodonts, square latest Carboniferous to Late Permian radiolarians and conodonts, and triangle LatePermian foraminifera, respectively.
and weakly metamorphosed olistostromal Fusaki rocks this belt (e.g. Miyashiro, 1973; Hirokawa, 1976; Hattori
& Shibata, 1982; Faure et al., 1988a; Banno & are thrust under the older and high- P/TTomuru
metamorphic rocks. The same features of metamorphic Nakajima, 1992). However, recent progress in geo-
chronology and petrology has led to a new interpret- geology and geochronology are observed in the rocks
of the Kuga-Nishiki area of western Chugoku (Fig. 3). ation (Nishimura et al., unpublished data).
The Nagasaki metamorphic rocks in the Nomo A summary of tectono–sedimentary and metamorphic
histories for both areas is shown in Fig. 4 (Ujiie & Peninsula (Fig. 1) can be classified into three geological
units, based on lithology, mineral paragenesis and Nishimura, 1992). The pair of the high- P/TTomuru
metamorphic rocks and the weakly metamorphosed K–Ar age. The three units are bounded by northwest-
dipping thrusts which collectively form a pile of olistostromal Fusaki rocks on the Ishigaki–Iriomote
Islands is comparable to that of the high- P/TSuo nappes: units A, B and C, from structurally higher to
lower levels (Fig. 5). metamorphic rocks and the adjacent Jurassic olisto-
stromal Kuga Group in the Kuga-Nishiki area. The Unit A consists mostly of basic and pelitic phyllites,
accompanied by metagabbroic rocks (metagabbro and southern Ryukyu Arc can, therefore, be interpreted as
the southwestern extension of the Inner Zone of amphibolite) dated at 590–460 Ma (Igi et al., 1979).
Both basic phyllites and metagabbroic rocks are Southwest Japan. The boundary thrust between the
Jurassic complex and the overlying pre-Jurassic characterized by the pumpellyite–actinolite assemblage.
This fact implies that both types of rocks underwent orogenic complex can be traced in both areas and
is named the Ishigaki-Kuga Tectonic Line (Isozaki the same metamorphism. K–Ar phengitic mica ages
from the pelitic phyllites are, in most cases, 200– & Nishimura, 1989). These results led Isozaki &
Nishimura (1989) and Ujiie & Nishimura (1992) to 150 Ma, and rarely as old as 254 and 248 Ma. Unit B
is composed of alternating beds of pelitic, psammitic propose a new tectonic framework for Southwest
Japan, as depicted in Fig. 3. and basic schists. These schists do not contain
porphyroblastic albite (albite spots), and so are denoted
‘non-spotted schists ’. The basic schists are metamor-Nagasaki metamorphic rocksphosed to epidote–actinolite and /or epidote–winchite
assemblages. The pelitic schists show chlorite zone A few small separate areas of high- P/Tmetamorphic
rocks, collectively named the Nagasaki metamorphic mineral assemblages. K–Ar phengitic mica ages from
the non-spotted pelitic schists indicate ages between belt, crop out on the western tip of Kyushu. Several
diVerent geotectonic settings have been proposed for 190–150 Ma. Late Cretaceous granite (90–80 Ma)
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GEOTECTONIC SUBDIVISION OF THE SANGUN BELT 133
Fig. 3. Comparison of the geotectonic framework of the Ryukyu Arc with Southwest Japan (modified from Isozaki & Nishimura,
1989). B.T.L.: Butsuzo Tectonic Line.
Fig. 4. Summary of tectono-sedimentary and metamorphic histories for pre-Tertiary rocks in the Ishigaki-Iriomote Islands and the
Kuga-Nishiki area (after Ujiie & Nishimura, 1992).
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134 Y. NISHIMURA
in the northeastern part of the Nomo Peninsula as
well as the Nishisonogi and Amakusa areas.
These facts suggest that the Nagasaki metamorphic
belt, previously regarded as a single metamorphic belt,
is in reality comprised of two metamorphic belts ofdiVerent ages and di Verent constituent rocks: units A
and B with 200–160 Ma ages (Triassic–Jurassic); and
unit C of c. 80 Ma (Late Cretaceous). It is therefore
reasonable to suggest that units A and B are assigned
to the Suo and /or Chizu belts in the Inner Zone, and
unit C to the Sanbagawa belt in the Outer Zone of
Southwest Japan. This assignment is strongly sup-
ported by the fact that Late Cretaceous granite of90–80 Ma intruded unit B but not unit C. It is highly
probable that the Wakimisaki-Fukabori thrust between
units A plus B and unit C (Fig. 5) corresponds to thepaleo-Median Tectonic Line as proposed by Isozaki
& Maruyama (1991).
Kiyama and Manotani metamorphic rocks
Several small metamorphic areas of unknown a Ynity,
such as the Kiyama and Manotani metamorphic rocks,
occur in central Kyushu (Fig. 1). Recent K–Ar dating
of these rocks have yielded new information concerningtheir tectonic a Yliation as follows.
Kabashima et al. (1995) re-examined the K–Ar
dating of the Kiyama high- P/Tmetamorphic rocks
and ascertained that eight phengitic mica ages of the
pelitic schists fall within the range of 306–290 Ma
(Early Carboniferous). The ages show a younger andFig. 5. Geological map of the Nomo Peninsula on the westernmore restricted range than those previously reported tip of Kyushu showing K–Ar ages from the Nagasaki
(343–303 Ma; Miller et al., 1963; Ueda & Onuki, 1968; metamorphic rocks and Cretaceous granite (after Nishimura
et al., unpublished data). Hyphenated numerals show biotite Ishizaka, 1972). It is therefore proposed that the
ages on the left and muscovite ages on the right. Non-Kiyama metamorphic rocks belong to the Sangun-hyphenated data represent phengitic mica ages. W-F thrust =Renge belt.Wakimisaki-Fukabori thrust.The Manotani metamorphic rocks, exposed 10 km
southeast of the Kiyama metamorphic rocks, are
polymetamorphic rocks, which had undergone a high- intruded schists of unit B and caused contact
metamorphism on this unit. K–Ar biotite and muscov- P/Tmetamorphism, and followed by the low- P/T
Ryoke metamorphism (Sato & Inoue, 1968; Karakida ite ages from the high-grade hornfelses give dates of
90–70 Ma. It is thus suspected that the younger ages et al., 1989). Okamoto et al. (1989) and Nagakawa
et al. (1993) reported five K–Ar ages ranging from (i.e. around 150 Ma) from the units A and B were
rejuvenated to some extent by the granitic intrusion. 214–168 Ma (Late Triassic–Middle Jurassic) for phen-
gitic micas from pelitic and psammitic schists. The Unit C is typified by a dominance of pelitic and
basic schists, which contain porphyroblastic albites, dates indicate that the Manotani metamorphic
rocks are correlative with those of the Suo and /or and these are therefore denoted ‘ spotted schists ’. The
unit is also accompanied by fairly large amounts of Chizu belts.meta–ophiolite, which consist of serpentinite and its
associated metabasic to metaacidic complex in theNishiki and Ikura areasform of a serpentinite me ´lange. In the basic metamor-
phic rocks, assemblages of epidote–winchite and epi- Although the Nishiki area west of Chugoku and the
Ikura area 160 km northeast of the Nishiki area have dote–barroisite are commonly recognized. However,
the assemblage of epidote–glaucophane is rare. The been classified as Suo and Chizu belts, respectively
(Fig. 1; Shibata & Nishimura, 1989; Nishimura, 1990), pelitic schists contain garnet and biotite (Miyazaki &
Nishiyama, 1989). K–Ar phengitic mica ages from the recent studies of these areas have provided new
information on the metamorphic history which requires pelitic spotted schists concentrate around 80 Ma. In
unit C there is no evidence of contact metamorphism. a re-consideration of this classification.
The Nishiki area is mainly occupied by the Suo Similar rocks to those of unit C are found throughout
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GEOTECTONIC SUBDIVISION OF THE SANGUN BELT 135
metamorphic rocks (Tsuno Group), which are overlain increasing degrees of graphitization (e.g. Itaya, 1981),
and supports the increase of metamorphic temperature by the weakly metamorphosed Permian accretionary
complex (Nishiki Group) consisting of clastic rocks of from zone A through zone B to zone C. K–Ar ages
were determined on 14 phengitic micas from the the Akiyoshi belt. The two groups are separated by
the Kitayama thrust. Palaeontological and sedimentol- samples which were used for obtaining the d002values.
They fell within a narrow range of 220 ±7 Ma. These ogical evidence from the Nishiki Group shows that
the age of deposition and accretion was Late Permian features, depicted in Fig. 6, clearly demonstrate that
all rocks of the Nishiki and Tsuno Groups underwent to earliest Triassic. In spite of the lack of fossil
evidence, the Tsuno Group may also have accumulated a single episode of high-pressure intermediate-type
metamorphism. in an ancient subduction zone during the Middle
Triassic, based on a K–Ar relict hornblende age Pre-Cretaceous metamorphic rocks in the Ikura area
can be classified into three units according to their (239 Ma) of an ophiolitic block (Nishimura &
Shibata, 1989). lithologies and metamorphic grades. These units are
separated from each other by north-dipping thrusts, The rocks of both groups have undergone regional
metamorphism and the metamorphic area is divided and collectively form a pile of nappes: the Ikura I,
Taniai and Oojaridani Formations, from structurally into the following three zones on the basis of the
mineral assemblages of basic rocks (Nishimura, 1971). higher to lower levels. The Ikura I Formation is
correlated with a greenstone–limestone unit of the Zone A is characterized by the pumpellyite–chlorite
assemblage and comprises the whole of the Nishiki Akiyoshi belt and the Oojaridani Formation with the
Chizu belt (Takeda & Nishimura, 1989). Group. The uppermost horizons of the Tsuno Group
belong to zone B, characterized by the pumpellyite– Studies on metamorphic mineral parageneses and
graphitization of carbonaceous material have revealed actinolite assemblage. Zone C is marked by epidote–
glaucophane and epidote–winchite assemblages and that the Ikura I rocks belong to the pumpellyite–
chlorite zone, the Taniai rocks to the pumpellyite– corresponds to the lower horizons of the Tsuno Group.
Zone boundaries are nearly parallel to bedding chlorite to lower pumpellyite–actinolite zones, and the
Oojaridani rocks to the pumpellyite–actinolite to schistosity, and the grade of metamorphism increases
downwards in the succession. epidote–glaucophane zones, and that these mineral
zones represent a progressive increase in metamorphic In order to further clarify the meaning of this
zonation, Nishimura et al. (1989) examined graphiti- temperature (Nishimura et al., 1996). K–Ar ages,
determined on 13 recrystallized phengitic micas from zation of carbonaceous material and K–Ar dates from
the pelitic metamorphic rocks along the Okuhata rocks of the three formations span 190–170 Ma (a
transition from the Early Middle to Jurassic), and River in the Nishiki area (Fig. 6). The apparent
interplanar spacing d002values of carbonaceous mate- clearly demonstrate that all rocks of the Ikura I, Taniai
and Oojaridani Formations underwent a single episode rial progressively decrease from the upper to the lower
horizons. The decrease in d002values correlates to of high-pressure intermediate metamorphism.
A summary of metamorphism and metamorphic
history along the Maebira-Oojaridani route in the
Ikura area is shown in Fig. 7 and compared with that
of the Nishiki area (Fig. 6). Unfortunately, the Ikura IFormation and the pumpellyite–actinolite zone of the
Oojaridani Formation are not exposed along this
route. The 190–170 Ma high- P/Tmetamorphism of
the Ikura area resembles the 220 Ma high- P/Tmeta-
morphism of the Nishiki area in that both of the
Sangun schists and rocks of the Akiyoshi belt under-went a similar type of metamorphism. The main
diVerence between them is the radiometric ages. These
features indicate that the Chizu and Suo eventspreviously defined by Nishimura (1990) can probably
be combined into one and the same subduction-
related orogen.
NEW GEOTECTONIC SUBDIVISIONFig. 6. d002data for carbonaceous material and K–Ar ages for
phengitic micas from pelitic metamorphic rocks along the
Two-fold division: Present model Okuhata River route in the Nishiki area (after Nishimura
et al., 1989). The horizontal scale represents distancesAll available radiometric ages reported from the measured normal to bedding schistosity of samples examined
and positions in zonal mapping. Sangun belt and its related areas are summarized in
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136 Y. NISHIMURA
Fig. 7. d002data for carbonaceous material and radiometric
ages for phengitic micas from pelitic metamorphic rocks along
the Maebira-Oojaridani route in the Ikura area (after
Nishimura et al., 1996). The horizontal scale represents
distances measured normal to bedding schistosity of samples
examined and positions in zonal mapping.
Fig. 8, along with the data cited in the previous
chapters. They spread from 362 to 153 Ma, from theearliest Carboniferous to latest Jurassic. Although
clearly rejuvenated and discordant ages discussed in
the previous chapters were excluded for the presentstudy, some ages obtained in areas where partial
rejuvenation by igneous intrusion is not obvious from
geological and petrographical studies may still have Fig. 8. Radiometric ages from high- P/Tschists in the Renge
been inadvertently included, especially in the Nomo and Suo belts including the Akiyoshi belt. Sources of data:
Miller et al. (1963), Hayase et al. (1968), Shibata & Nozawa Peninsula.
(1968), Shibata et al. (1968, 1970, 1972, 1979, 1984), Ueda &Excluding the rejuvenated and discordant ages suchOnuki (1968), Shibata & Igi (1969), Ishizaka (1972), Shibataas around 150 Ma in Nomo, 362 Ma in the Hida& Ito (1978), Matsumoto et al. (1983), Nishimura (1984),marginal belt, and 343 Ma in Kiyama, the radiometric Faure et al. (1988b), Fukudomi et al. (1989), Nishimura et al.
(1989, 1996, unpublished data), Okamoto et al. (1989), Shibata ages shown in Fig. 8 can be divided into two age
& Nishimura (1989), Ujiie & Nishimura (1992), Nagakawa groups: the older group is 330–280 Ma (Carboniferous–et al. (1993), and Kabashima et al. (1995).Permian) and the younger group is 230–160 Ma
(Triassic–Jurassic). The author proposed naming the
older group the Renge belt, and the younger group Maruyama, 1991); the Akiyoshi belt is a Late Permian
(c. 250 Ma) accretionary complex including 230– the Suo belt (Figs 8 & 9). The former includes the
Sangun-Renge belt defined by Nishimura (1990) and 160 Ma low-grade regional metamorphic rocks; the
Maizuru belt is a Middle Permian accretionary the Kiyama area in central Kyushu, and the latter
includes the Suo and Chizu belts of Nishimura (1990), complex with the Yakuno ophiolitic complex; the
ultra-Tamba belt represents a Early Triassic–Late the Nomo (units A & B) and Manotani areas in
central Kyushu, and the Ishigaki-Iriomote Islands in Permian (?) accretionary complex; the Mino-Tamba
belt is a Jurassic accretionary complex, which grades the southern Ryukyu Arc.
The Renge and Suo belts occur as gently north- southwards into the Ryoke metamorphic rocks.
A two-fold division of the Sangun belt into the dipping subhorizontal nappes, bounded by low-angle
faults and thrusts. It is thus considered that the Inner Renge and Suo belts is geochronologically and geotec-
tonically preferable to the three-fold division previously Zone of Southwest Japan is composed of, from north
to south, the Hida-Oki, Renge, Akiyoshi, Suo, proposed. The following sections will briefly describe
the distinctive features of the two belts (Fig. 9). Maizuru plus ultra-Tamba, Mino-Tamba, and Ryoke
belts, as shown in Fig. 9. The characteristics of these
belts (except for the Renge and Suo belts) are givenRenge belt: 330–280 Ma (Carboniferous–Permian) groupin the following: the Hida-Oki belt consists of a
2.0 Ga–180 Ma gneiss and granite complex probably This is the oldest high- P/Tmetamorphic belt in the
Japanese Islands (330–280 Ma), and extends from derived from a part of the Yangtze craton (Isozaki &
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GEOTECTONIC SUBDIVISION OF THE SANGUN BELT 137
340 Ma, i.e. Ordovician to Carboniferous periods, as
shown in Fig. 10 (Nishimura & Shibata, 1989).
The Renge rocks are considered to be in fault
and thrust contact with adjacent units, and to
show collectively a subhorizontal nappe. Isozaki &Maruyama (1991), in their new geotectonic subdivision
of the Japanese Islands, propose that the upper
boundary of the Renge belt is formed by the Nagato-
Hida Tectonic Line, overlain by the 580–450 Ma
ophiolite of the O-eyama belt, whereas the lowerboundary is represented by an as yet unconfirmed
thrust separating it from the underlying Akiyoshi belt.
In the Nagato tectonic zone, Kabashima et al. (1993)
found the boundary thrust between the Renge schists
above and the Permian accretionary complex of the
Akiyoshi belt below. However, the scanty evidence
Fig. 9. New geotectonic subdivision and areal extent of theleaves room for re-examination of the conception ofSangun belt (Renge and Suo belts) along with the geotectonicthe O-eyama belt itself as well as the geotectonicframework in the Inner Zone of Southwest Japan. Therelationship between the O-eyama belt and the Renge geotectonic subdivision of the Ryukyu Arc and its correlation
are shown separately in Fig. 3. belt.
Suo belt: 230–160 Ma (Triassic–Jurassic) groupnorthern Kyushu through the San-in coastal regions
to the Hida marginal belt, which lies on the south of This belt is characterized by 230–160 Ma high- P/T
schists (Suo schists) closely associated with a weakly the Hida-Oki belt (Fig. 9). Outcrops belonging to this
belt are narrow and discontinuous, being found from metamorphosed Permian accretionary complex of the
Akiyoshi belt. Outcrops belonging to this belt are west to east in the Wakamiya area, the Toyoga-dake
area in the Nagato tectonic zone, the Shitani Formation widespread in west Kinki to north and central Kyushu
via Chugoku, and are found to lie in a NE-SW to of the Wakasa area (Fig. 1), and several small areas in
the Hida marginal belt (Komatsu, 1990). The Kiyama ENE-WSW trend (Fig. 1). They extend farther to the
southwest into the Nomo, Manotani and Ishigaki- area in central Kyushu (Figs 1 & 9) is a southern
extension of the belt. The eastern extension of the Iriomote areas (Figs 9 & 3). The region comprising
these areas is here named the Suo belt, which combines Hida marginal belt (Renge belt in this paper) has long
been considered to be the Tanigawa-dake zone in the the Suo and the Chizu belts defined by Nishimura
(1990). ‘Suo’ is the old local name of eastern Yamaguchi Joetsu belt on the eastern side of the Itoigawa-
Shizuoka Tectonic Line. However, there is an alterna- Prefecture, where typical high- P/Tschists and their
related weakly metamorphosed rocks of the Akiyoshi tive opinion that the zone is an eastern extension of
the Suo belt (Komatsu, 1990). As no radiometric data belt are distributed.
The metamorphic facies and facies series of the belt have been reported, this problem remains unsolved.
Although this belt is nearly identical with the are of the prehnite–pumpellyite to lower pumpellyite–
actinolite facies in the Akiyoshi rocks, and of the Sangun-Renge belt defined by Nishimura (1990), the
Renge belt as a new name is used for this belt pumpellyite–actinolite to glaucophane–schist facies
with rare epidode–amphibolite facies in the Suo henceforth. ‘Renge’ is derived from the name of a
mountain in the Hida Mountains and was originally schists (Hashimoto, 1972; Nishimura et al., 1977).
Serpentinites and related metagabbroic rocks sporadi- used for the type area name of the Renge metamorphic
rocks in the Hida marginal belt (Ishii, 1937). cally occur as small blocks or lenses throughout the
belt. These are interpreted as parts of a dismembered The rocks of the Renge belt are composed of high-
P/Tschists metamorphosed mainly at the glauco- ophiolite. The metagabbroic rocks including amphibo-
lite underwent initial ocean-floor metamorphism in the phane–schist to epidote–amphibolite facies, the proto-
liths of which are an accretionary complex of unknown epidote–amphibolite to amphibolite facies, and sub-
sequent Suo metamorphism in the pumpellyite–actino- age, the main protoliths being mudstone, basaltic rocks
and chert. The belt is also characterized by substantial lite to glaucophane–schist facies. K–Ar ages of relict
hornblendes are 590 to 220 Ma, i.e. Precambrian to amounts of serpentinite, which sometimes contains
blocks of metagabbro, metagranite, eclogite and jadeit- Triassic time, as shown in Fig. 10 (Nishimura &
Shibata, 1989; Igi et al., 1979). ite. All these rocks appear to reflect portions of a
dismembered ophiolite or serpentinite me ´lange. The The Suo belt is a north-dipping nappe sandwiched
between the Akiyoshi belt above and the Mino-Tamba metagabbro and metagranite that may have undergone
ocean-floor metamorphism in the epidote–amphibolite or Maizuru belts below (Fig. 9). The boundaries of
this belt have been found in many areas: the former, to amphibolite facies, have radiometric ages of 470–
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138 Y. NISHIMURA
Fig. 10. Radiometric ages from metagabbroic rocks in the Renge and Suo belts showing metamorphic age groups of high- P/T
schists in Figs 8 & 9. Sources of data: Kawano et al. (1966), Hayase & Ishizaka (1967), Shibata et al. (1972, 1977, 1979, 1980,
1984), Ishizaka & Yanagi (1975), Shibata & Murakami (1975), Murakami et al. (1977), Murakami & Nishimura (1979), Igi et al.
(1979), Matsumoto et al. (1983), and Nishimura & Shibata (1989).
for example, the Kitayama thrust in the Nishiki area (2) The Suo belt is characterized by 2 30–160 Ma
(Triassic–Jurassic) high- P/Tschists closely associated (Nishimura et al., 1989) and the Harada thrust in the
Ikura area (Nishimura et al., 1996); and the latter, for with a weakly metamorphosed Permian accretionary
complex of the Akiyoshi belt. Outcrops belonging to example, the Notanigawa fault in the Kuga area
(Toyohara, 1974) and the Sokobaru thrust in Ishigaki the belt are widespread in west Kinki to north and
central Kyushu via Chugoku, extending farther to the Island (Fig. 2; Fujii & Kizaki, 1983).
Protolith assemblages of the Suo schists are southwest into the Nomo, Manotani and Ishigaki-
Iriomote areas (Figs 8 & 9). The Akiyoshi rocks classified into two types. The first consists predomi-
nantly of greenstone with significant amounts of underwent metamorphism in the prehnite–pumpellyite
and lower pumpellyite–actinolite facies. In contrast, mudstone and sandstone but with only scarce chert.
The second assemblage is predominantly bedded chert the Suo schists underwent metamorphism in the
pumpellyite–actinolite, glaucophane–schist, and epi- but also contains clastic and basic rocks. The first
assemblage occurs throughout this belt, whereas the dote–amphibolite facies. The remains of meta–ophiol-
ites occur sporadically as small blocks or lenses second is confined to the eastern areas of the belt (i.e.
Wakasa, Katsuyama, Tsukita and Masuda areas in throughout the belt and relict hornblendes have K–Ar
ages of 590–220 Ma (Fig. 10). Metamorphic rocks Fig. 1). Metamorphic rocks derived from the chert-
predominant protolith may, in some or most cases, derived from the chert-predominant protolith may, in
some or most cases, represent a higher metamorphic represent a higher grade part of the Jurassic
accretionary complex in the Mino-Tamba belt grade part of the Jurassic accretionary complex of the
Mino-Tamba belt. (Hayasaka, 1987; Nishimura, 1990). However, this
suggestion requires more detailed study to be verified. (3) Each of the two belts (Renge and Suo) is a
north-dipping subhorizontal nappe, bounded aboveand below by low-angle faults and thrusts. TheCONCLUSIONSgeotectonic framework in the Inner Zone of Southwest
Japan is made up of, from north to south, the Hida- The three-fold division proposed by Nishimura (1990)
and the areal extent of the Sangun (metamorphic) belt Oki, Renge, Akiyoshi, Suo, Maizuru plus ultra-Tamba,
Mino-Tamba, and Ryoke belts with a tectonically were re-examined on the basis of recent studies of
metamorphic geology, tectonics and geochronology. A downward-younging polarity (Fig. 9). This has resulted
from stepwise accretions during Palaeozoic to new two-fold division of the Sangun belt into the
Renge and the Suo belts is proposed based on di Verent Mesozoic time.
ages and contrasting conditions of formation. Thefollowing conclusions are reached:
(1) The Renge belt is the oldest of the high- P/TACKNOWLEDGEMENTSmetamorphic belts in the Japanese Islands and extends
from northern Kyushu including the Kiyama area, The author is indebted to Professor S. Banno, Dr S. R.
Wallis and reviewers, Dr Y. Kawachi and Professor through the San-in coastal regions, to the Hida marginal
belt, which lies on the south of the Hida-Oki belt W. G. Ernst, for improving the manuscript, both
scientifically and linguistically, and to Dr Y. Isozaki (Fig. 9). It is typified by 330–280 Ma (Carboniferous–
Permian) schists metamorphosed in the glaucophane– for discussion and constructive comments. This study
was supported by the Grant-in-Aid from the Japan schist to epidote–amphibolite facies, and is accompanied
by meta–ophiolite (470–340 Ma) in the form of serpentin- Ministry of Education, Culture and Science
(No. 07454121). ite me´lange (Figs 8 & 10).
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GEOTECTONIC SUBDIVISION OF THE SANGUN BELT 139
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Journal Metamorphic Geology - 2004 - NISHIMURA - Geotectonic subdivision and areal extent of the Sangun belt Inner Zone of.txt
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ELSEVIER Tectonophysics 284 (1998) 135-150 TECTONOPHYSICS
History of basin formation and tectonic evolution at the termination of
a large transcurrent fault system: deformation mode of central Kyushu,
Japan
Yasuto Itoh a,*, Keiji Takemura b, Hiroki Kamata c
"Department of Earth Sciences, CIAS, Osaka Prefecture Universi~., Gakuencho 1-1, Sakai, Osaka 599, Japan
h Department of Geophysics, Kyoto Universit3, Sakyo-ku, Kyoto 606-01, Japan
c Geological Survey' of Japan, 4-1-67 Otemae, Chuo-ku, Osaka 540, Japan
Received 15 October 1996; accepted 19 June 1997
Abstract
The central part of the Kyushu Island is a locus of active arc-volcanism associated with subduction of the Philippine
Sea plate. Being located at a bend in the boundary between the Philippine Sea and Eurasian plates, a complicated lateral
displacement has generated half-grabens, a zone of strike-slip, and rhomboidal basins filled with volcanic material in
central Kyushu. Initial northward subduction in latest Miocene time activated the N-S-trending Kokura-Tagawa Tectonic
Line (KTL) as a left-lateral slip fault that bounds the western margin of the area of volcano-tectonic depression, the Hohi
volcanic zone, which was initially formed about 6 Ma ago. The depocenter, which is now called Kuju basin, in the Hohi
volcanic zone was at the corner of the KTL and the E-W-trending Median Tectonic Line (MTL). Relative convergence
direction of the Philippine Sea plate shifted counterclockwise at about 1.5 Ma, and active motion on the KTL declined. At
the same time, west-northwestward subduction enhanced right slip on the MTL. Geologic, gravimetric, and seismic data
indicate that the MTL has shifted its active trace northward in central Kyushu as far as 10 km. As a result, the depocenter
adjacent to the transcurrent fault migrated northeastward in the Hohi volcanic zone, specifically, from Kuju basin (Pliocene)
via Shonai basin (early Quaternary) to Beppu Bay basin (late Quaternary). The latest depocenter of the Beppu Bay is
surrounded by active faults that clearly delineate a rhomboidal basin on the MTL. Central Kyushu exemplifies the basin
forming history and tectonic evolution at the termination of a large transcurrent fault system. © 1998 Elsevier Science
B.V. All rights reserved.
Keywords: rhomboidal basin; Kyushu; flower structure; transcurrent fault; Philippine Sea plate; Median Tectonic Line
I. Introduction
Kyushu Island lies at the junction of the Ryukyu
and the Southwest Japan Arcs (Fig. 1). Arrangement
of pre-Neogene terranes within Kyushu suggests that
the island is basically the westernmost part of a co-
* Corresponding author. Fax: -+-81 722 55 2981. herent block of southwest Japan, although a Miocene
bending deformation is assumed because of a fan-
shaped back-arc spreading of the Sea of Japan at
around 15 Ma (e.g., Otofuji et al., 1991). Since
latest Miocene time, Kyushu has been under the
influence of subduction of the Philippine Sea plate
(Kamata and Kodama, 1994). Major back-arc folds
caused by N-S shortening deformation in south-
0040-1951/98/$19.00 © 1998 Elsevier Science B.V. All rights reserved.
PII S0040-1951(97)00167-4
136 Y. ltoh et al./Tectonophysics 284 (1998) 135-150
! i iiiiiiiiiil I i iil!i
ii ii:~iii~iiiiiiiiiiii? ~
..... i~:i ;ii;i~ !i~i~ i I !! .....
Q
Fig. 1. Index map around the Japanese Islands (ISTL = Itoigawa-Shizuoka Tectonic Line). The recent relative convergence vector and its
convergent rate of the Philippine Sea plate is shown by open arrow (Seno, 1977). The darkened arrow is the convergence vector inferred
from the latest Miocene tectonic features in southwest Japan (Itoh et al., 1992; Yamamoto, 1993).
west Japan during late Miocene time (Itoh et al.,
1992; Yamamoto, 1993) chronicle the initial stage
of subduction which activated the transcurrent N-
S Kokura-Tagawa Tectonic Line (KTL) (Fig. 2).
Due to a counterclockwise shift of convergence di-
rection at about 1-1.5 Ma (Fig. 1; Seno, 1985;
Nakamura et al., 1987; Yamazaki and Okamura,
1989), the Philippine Sea plate activated the tran-
scurrent E-W Median Tectonic Line (MTL) (Fig. 2)
at the concave bend of the overriding Eurasian
plate. The KTL and MTL bound the western and south-
eastem margins of the Hohi volcanic zone, respec-
tively (Figs. 2 and 3; Kamata, 1992). The Hohi
volcanic zone is a composite depression character-
ized by a conspicuous negative box-shaped Bouguer
anomaly and buried magmatic material of >5000
km 3 (Fig. 4; Kamata, 1989b). In central Kyushu,
some assume an E-W-trending rift zone from the
Beppu Bay to the Shimabara Peninsula (Fig. 2).
Based on repeated triangulation and trilateration sur-
veys during the last 90 years, Tada (1985) argued
Y. Itoh et al./Tectonophysics 284 (1998) 135-150 137
!~ 34N
i T ..........
130E 131E 132E
i~3N
I
N
~hhl~om~ Pen,
• ~l~olt & dilq~caaem i
0 I00 km ~"
[ I i I dis~ transcurrent fault
Fig. 2. Index of Kyushu Island (KTL = Kokura-Tagawa Tectonic Line; MTL = Median Tectonic Line). The star denotes the site of a
drilling study (Noi and Ito, 1991). The inset shows a two-dimensional slip-line field (thin lines) in front of an indenter (shaded triangle)
after Itoh and Takemura (1994) with crustal displacement vectors (arrows) after Tada (1985).
that the extension basins in central Kyushu formed
by a N-S breakup of the continental crust. His es-
timate for the extension rate of 14 mm/year around
the Shimabara Peninsula is, however, too large to
accommodate the real separation around the Hohi
volcanic zone. If the volcanic depression extended at
the constant rate proposed by Tada, then the north-
ern and southern flanks should have merged at 3
Ma, an age which is contradictory to the geologic observation that Pliocene volcanic rocks are present
within this depression (Kamata, 1989b). Thus, the
geodetic data may be attributed to a local extension
around the Shimabara Peninsula, where Pliocene
basalts have the chemical signature of oceanic is-
land rocks (Nakada and Kamata, 1988). Kamata and
Kodama (1994) proposed that oblique subduction of
the Philippine Sea plate beneath the Southwest Japan
Arc detached a fore-arc sliver that was displaced
138 Y ltoh et al./Tectonophysics 284 (1998) 135-150
STUDY AREA Suo-nada Sea
• O • • • • • • • • • • • • • •
• Hohi VolcanicZone ~.
• •o
•
~Aso Caldera ~_~ 10 20km
131°E Beppu Bay Pliocene - Pleistocene
volcanics and sediments
Miocene volcanics and ! I
intrusives
Pre-Cenozoic basement
(generalized)
Fault ]
g
y 1320E .33oN
Fig, 3, Geologic map of central Kyushu modified from Kamata (1989b) and Kamp and Takemura (1993). Dots indicate the outline of the
Hohi volcanic zone.
along the MTL and hence formed the Hohi volcanic
zone by pull-apart motion.
An alternative interpretation for the N-S exten-
sion in Kyushu was proposed from the viewpoint of
neotectonic framework. From offshore active fault
distribution and seismological data, Kimura (1996)
stated that Kyushu is thrusting upon the fore-arc
sliver of southwest Japan. Itoh and Takemura (1994)
suggested that the distribution of active transcurrent
faults in Kyushu follows a slip-line field caused by
westward indentation of the fore-arc sliver associated
with the Quaternary dextral motions on the MTL
(see inset of Fig. 2). This framework also accords
with prevailing extension in the western back-arc region of Kyushu. Results of a seismic survey in the
Amakusa-nada Sea (Fig. 2) delineates horst-graben
features developed since the Pliocene time (Ishiwada
et al., 1992). A drilling study near the Shimabara
Peninsula demonstrates that the thick Quaternary de-
posits rest directly on Paleogene deposits (Noi and
It•, 1991; star in Fig, 2). Therefore, the crust of
Kyushu is not simply separated in N-S direction, but
extruded by the westward indentation of an adjacent
landmass, fore-arc sliver of southwest Japan. The
inset of Fig. 2 demonstrates that the geodetic data by
Tada (1985) rather accord with the extrusion theory.
The right-lateral MTL activity caused not only
crustal extrusion, but also a series of basins to form
Y ltoh et al./Tectonophysics 284 (1998) 135-150 139
1310E 132°E
,33°N
Fig. 4. Bouguer anomaly map and nomenclature of depressions. The anomaly map is compiled from Komazawa and Kamata (1985) and
Takemura et al. (1993). Abbreviations for basins are: BB = Beppu Bay, SN = Shonai, KJ = Kuju, SY : Shin-Yabakei, SS = Shishimuta,
and AS = Aso. The shaded zone is a gravity slope coincident with the western margin of the Hohi volcanic zone (Kamata, 1992).
in Kyushu. Takemura et al. (1994) recognized on
seismic profiles a basin centered at the present Beppu
Bay (Figs. 3 and 4), bounded by a rollover master
fault and transcurrent flower structures constituting
rhombic margins. The Beppu Bay basin is the east-
ernmost and youngest depocenter in the composite
depression of the Hohi volcanic zone which has ex-
isted since latest Miocene time (ca. 6 Ma) (Kamata,
1993). In this article, we describe the history of basin
formation and tectonic evolution of the Hohi vol- canic zone in central Kyushu, at the termination of a
large transcurrent fault system.
2. Geologic setting
2.1. Basement geology
The basement of the Neogene basins in cen-
tral Kyushu (Fig. 3) consists of Cretaceous granite,
marine sedimentary rocks, and Paleozoic schist dis-
140 Y ltoh et al. / Tectonophysics 284 (1998) 135-150
tributed around the Hohi volcanic zone. Geothermal
wells in the Hohi volcanic zone encountered granite
and low-pressure metamorphic rocks of Cretaceous
age at depths of more than 1000 m (Sasada, 1987),
which resemble the granitic basement of the Ryoke
terrane exposed north of the volcanic zone. South of
the Hohi volcanic zone, high-pressure metamorphic
rocks are distributed and assigned to the Sambagawa
terrane of the Outer Zone of southwest Japan. Thus
the MTL, defined as the boundary between the pre-
Neogene Ryoke and Sambagawa terranes, coincides
with the south margin of the Hohi volcanic zone
(Kamata, 1992). This part of the MTL has been inac-
tive since Cretaceous time (Ichikawa, 1980), except
for the Pliocene to Quaternary dextral displacement
which is discussed later.
2.2. Late Cenozoic basins
A remarkable negative Bouguer anomaly char-
acterizes the Hohi depression and clearly delineates
several depocenters (Fig. 4). The elongate Kuju basin (KJ) is situated near the southwest corner of the Hohi
volcanic zone. The N-S-trending basement ridge of
the Mizuwake-Toge divides the Shonai basin (SN)
and the Shishimuta negative anomaly (SS) that was
interpreted as an early Pleistocene buried Shishimuta
caldera (Kamata, 1989a). Beppu Bay (BB) is the
easternmost and largest basin in the study area, and
has been uniquely active since middle Pleistocene
time.
2.3. Volcanism in the Hohi volcanic zone
The relief on the basement of the Hohi volcanic
zone is covered by extensive Pliocene and Pleis-
tocene volcanic rocks. The volcanism produced an-
desite to rhyolite lavas, large-scale pyroclastic-flow
deposits, and very little basaltic rocks. Their radio-
metric ages range from 6 Ma to recent (Kamata,
1989b). From the viewpoint of geochemistry, the
K20 content increases with time (Table l; Kamata,
1989b) and older calc-alkalic activity (>3 Ma) was
followed by eruption of tholeiitic rocks (Nakada and
Table 1
A synthesis of evolution of the Hohi volcanic zone
Geologic period PHS Strain in Hold Volcanism in Hohi Basin Fault motion
subduction volcanic zone volcanic zone subsidence
western part: ~ ~ ~_ ~,
! o~ Middle Pleisto_ , .... contraction i~ ~ ~ ~'~ BB
I'~ cene - Holocene wl, w eastern part: .~ j:~ .~ ~. MTL2 ++ (R) +~ extension ~ .~ ~ rhombic
-Episode 3 (0. 7 Ma): Shift of active trace of MTL A ~ ~ A
southwestern part: i i ~ i { • : J o Early WNW contraction i ', -- i SN, BB
~ Pleistocene northeastern part: i i A rhombic MTL1 ++ (R)
!extension ~ i i
-Episode 2 (1.5 Ma): Counter-clockwise shift of Philippine Sea Plate (PHS) subduction
I
it KJ, SN, BB KTL + (L) NNW-SSE I i ! :
• ~ plE~e~Yocene N-NNW extension i
I (regional regime) ! ! :, i asymmetricIMTL0 + (R)
o
o!~ KJ, SN, BBI ICIT, ++ (L) ..... NNW-SSE , ~ = Late Miocene - • ~ .
iocene N- NNW extension [ ~ ~ ~.~ P1 (regional regime) _~ ~ ~ asymmetri~MTL0 + (R) 1
e • o p.
~--=~'+" ~ (6 Ma)" ,..i~.~.... <.+ u..J.i .,,,I,.o.,,.
See Fig. 4 for abbreviations of basins. ++ = high fault activity; + : low fault activity; (R) = right-lateral motion; (1) = left-lateral
motion.
Y ltoh et al./Tectonophysics 284 (1998) 135-150 141
Kamata, 1991). The style of emplacement changed
from the Pliocene fissure eruptions without sig-
nificant pyroclastic flow to the Pleistocene spo-
radic activity forming groups of monogenetic volca-
noes and small-scale stratovolcanoes (Kamata et al.,
1988).
2.4. Borehole data
Core samples in the Pliocene-Pleistocene vol-
canic pile are dominated by andesitic lava flows and
breccia. Minor intercalations of sedimentary deposits
indicate continuous volcanism from the beginning of
the Hohi volcanic zone. The K-Ar ages of core vol-
canic rocks are generally older toward the bottom
of the graben, reaching a maximum of at least 4
Ma (Fig. 5; Kamata, 1989b). The volcanic activity,
which began at about 6 Ma, became concentrated
with time toward the center of the graben, burying
successively erupted material (Kamata, 1989b).
The volcaniclastic rocks filling the Shonai basin
(SN, Fig. 4) are assigned to the Quaternary (Hase
and Iwauchi, 1993; NEDO, 1991). Most of the basin
is underlain by the Kawanishi andesite (1.4-1.6 Ma)
(Hoshizumi and Kamata, 1991), the Shikido pyro-
clastic flow deposits erupted at about 1.3-1.5 Ma
(Takemura et al., 1988), the Yabakei pyroclastic flow
erupted at 1 Ma (Kamata, 1989a), and overlain by
the Yufugawa pyroclastic flow erupted at 0.6 Ma
(Hoshizumi and Kamata, 1991; Takemura and Dan-
hara, 1993). Thus, the Shonai basin is a young
depocenter when compared to the Kuju basin within
the Hohi volcanic zone.
2.5. Active faults
The Hohi volcanic zone is studded with a num-
ber of E-W-striking normal faults (Fig. 6). West of
the Shonai basin, faults are arranged in en echelon
pattern implying rhomboidal basin formation. The
northern slope of the Beppu Bay is an active fault,
the Karakiyama fault (Fig. 6), which has a clear
right-lateral displacement (Research Group for Ac-
tive Faults, 1991). Takemura et al. (1994) recognized
that north side of an active rhomboidal basin coin-
cides with the Karakiyama and Kanagoe faults. A
NW-SE-trending listric fault bounds the east of a
hanging-wall collapse (expressed as "rollover") and an ENE-WSW transcurrent fault in the Beppu Bay
(expressed as "flower structure") defines south side
of the basin. The western side of the rhomboid is the
Asamigawa fault with 285 m of cumulative vertical
separation (Takemura and Danhara, 1993).
3. Discussion
Quaternary activity of the MTL has been stud-
ied on southwest Japan (Tsutsumi et al., 1991), and
some right-lateral motions are identified. Ito et al.
(1996) interpreted reflection seismic data across the
MTL on Shikoku Island and found that pre-Tertiary
terranes have low-angle north-dipping contact, al-
though an active trace of the MTL is hard to detect
from their data. Yamakita et al. (1995) reported a
low-angle oblique thrust along the Sashu fault in the
early Oligocene, to the south of the Beppu Bay, as
an activity of Paleo-MTL. Shimazaki et al. (1986)
studied single-channel seismic and boring data in
the Beppu Bay and estimated recurrence interval of
the active trace of MTL from vertical fault displace-
ments. However, listric faults in shallow depths in
their work are a subordinate structure within a large
rhomboidal basin in the bay (Takemura et al., 1994).
Thus, the westernmost segment of the MTL around
Kyushu has not been thoroughly studied from the
viewpoint of neotectonics.
3.1. Shift of the active trace of the MTL
Seismic reflection and gravity data of the Beppu
Bay (Takemura et al., 1994) indicate that the base-
ment structure is a half-graben, which closely re-
sembles the onland area of the Hohi volcanic zone
(Fig. 5); in both areas pre-Tertiary rocks dip asym-
metrically southward (Kamata, 1993). Because the
Beppu Bay has not been the source vent of a large-
scale pyroclastic flow, its negative anomaly is un-
likely to originate from volcanic depression such
as a caldera. Flower structures running through the
Beppu Bay led Takemura et al. (1994) to interpret
the structure accompanying dextral movements as
the westernmost part of the active MTL. There are
two major developments of flower structures on both
flanks of buried basement depression of the bay
(Fig. 7). One on the southern flank (FSI) seems
to be overlain by intact late Quaternary sediments,
142 Y. ltoh et al./Tectonophysics 284 (1998) 135-150
£'AO u~o001 O)
"C
o.,~o,*-.
~'o* ~-80
"~"-"!t, "., .," ...... -- "'" ~z " i!~" ""
'. L.... ,.-'.,
" i !: 15 " ¢" • .... • ........................... ,,.. ~ '~
:";!. j""" "" ""'k.. / .................. ~ .... ," :"
~ ~,'
, =r" I¢0 ~-"'"* ~:", .......
rSVp" , • ,. ,~ . -"
""'" S" o" 0 '*"*,...
";> '."' E = ,: ~, ~
,..:" !
z
o*} ',.. ........... ?... ) "......../
0 .=.
.,,,..,
.o
0
0 ,=~
0
r=,
F_,
Z
<~
o
,l) "~
0 oJ
o
E.)
Y. ltoh et al. / Tectonophysics 284 (1998) 135-150
Z
m
"=~ <~
e-
~.~
~.~ r~
~,'~
e~
~ 0
~°
~ "0
"0
e- e-,
• =~ ~ 143
144 Y. ltoh et al./Tectonophysics 284 (1998) 135-150
~t
0
(m) 4~d~43 i
E C)
e-
8
0
.o
t~
.o
g.
r~
¢q
o~
e-.
c)
.o
,,.m
2
t~
o
....q
Y. Itoh et al./Tectonophysics 284 (1998) 135-150 145
whereas the other on the northern flank (FS2) cuts
the entire sedimentary sequence within the depres-
sion. The southern structural trace is approximately
under Oita City, and the Tokiyama fault in the south
of Shonai basin (see Fig. 6) is a possible extension of
the fault. The fault trace has almost been buried by
the recent sediment and soil. It is suggested therefore
that the southern structure (FS1) is an abandoned
trace of the MTL, the role of which was taken
over by the latest northern structure (FS2) extending
along the southern shoreline of the Beppu Bay.
Horizon A (Fig. 7) is drawn on the basis of a
sharp contrast of reflection pattern on the seismic
profile. Top of the FS 1 traces roughly coincide with
the horizon. The sedimentary thickness of the lower
unit suggests that the depocenter of the Beppu Bay
was on the fossil MTL trace of the FS 1 before late
Pleistocene. On the other hand, the depocenter of the
basin during late Pleistocene and Holocene time is
assigned along the FS2, based on the thickness of the
upper unit and abnormally rapid sedimentation along
the southern shoreline of the bay (Noi, 1987).
Another possible trace of the transcurrent fault
lies south of the above two traces in the Beppu Bay.
Yoshioka (1992) pointed out that the Sekinan Group
with ages of late Pliocene and early Pleistocene
had been deformed by right-lateral displacements on
a fault between the Neogene basin and basement
mountains (MTL in Fig. 6). This is the fault de-
scribed, in the section about the geologic setting, as
the inactive trace of the MTL. That inactive trace
also coincides with the southern margin of the Hohi
volcanic zone (Kamata, 1992). The amount of north-
ward migration of the MTL traces thus exceeds 10
km around the Beppu Bay.
3.2. Initiation of the Hohi volcanic zone
Initial subduction of the Philippine Sea plate was
in northern direction in the latest Miocene time
(Fig. 1; Itoh et al., 1992; Yamamoto, 1993). Itoh and
Nagasaki (1996) showed, using seismic and drilling
data, that the northern back-arc of southwest Japan
and Kyushu is strongly deformed during the late
Miocene. The Philippine Sea plate resumed subduc-
tion at about 6 Ma after a halt or slow down of
subduction during the period 10-6 Ma (Kamata and
Kodama, 1994). It is therefore probable that the N- S-trending KTL was first activated as a left-lateral
fault at the bend in the plate boundary, whose mo-
tion was accommodated by formation of E-W half-
graben on the pre-existing crustal break, the MTL.
We interpret that this coincides with the initiation of
the Hohi volcanic zone (Episode 1 in Table 1). A
N-S cross section of the Hohi volcanic zone (line
c-d in Fig. 5) shows that the depression still has an
asymmetric shape that deepens southward.
In Fig. 8, azimuths of dike swarms and veins in
Au and Sb mines are plotted with the distribution of
the Pliocene volcanic rocks. Because the Au and Sb
deposits are closely related to the Pliocene volcanism
which induced vein-type mineralization (Kamata,
1989b), the E-W trend of the data indicates a highly
uniform ~rHmax direction in the Hohi volcanic zone
during Pliocene time (Table 1). A similar stress
direction was reported for the SS area now covered
by the Pleistocene volcanic rocks around the buried
Shishimuta caldera. Kamata et al. (1988) reported
that a local, weak compressional stress field initiated
in the middle Pleistocene (-0.7 Ma) and started to
overlap the previously dominant N-S tensional stress
field (Table 1).
Although vigorous motion on the KTL might
have declined by the end of Miocene time, consider-
ing structural features of the early Pliocene volcanic
rocks of Mt. Hikosan (Fig. 9a; Hikosan Collaborative
Research Group, 1992) on the fault trace, the uni-
form E-W tYHmax mentioned before implies that the
formation of a graben was active through Pliocene
and early Pleistocene time (Table 1).
As for the location of the western side of the
Hohi volcanic zone, Kamata (1992) stated that the
flank of a negative gravity zone between Hita and
Miyanoharu (shaded zone in Fig. 4) is a probable
candidate. The zone coincides with the southern ex-
tension of the KTL (Fig. 9a). East of the proposed
boundary, however, there are exposures of Creta-
ceous basement (granite and low-pressure metamor-
phic rock) around Kashinomure (Fig. 5), and bore-
holes on the west of Shishimuta (SS) depression en-
countered identical basement rocks at depths around
1000 m (DW-7 and DY-2 in Fig. 5; Sasada, 1987).
Therefore, we propose that the Pliocene northward
motion of the Philippine Sea plate activated a left-
lateral slip zone between the KTL and the N-S
lineament connecting the western sides of the Kuju
146 Y. Itoh et al. / Tectonophysics 284 (1998) 135-150
N
I ~l-l~ aTsurumi
P _"-~¢~-~Y,~- ss SN A-nTIIIIIg
[ [ [ ~'k_ ...47 a~K/[[[[[ [1[ ~ Lava d°mes y°unger than O' I Ma
/~'~ ~ a~lll~llll !~1111111 F=I Mi°~ v°l~i~ =d i~ves
~-~ ~lllll~l~dll I]TI~ Pre-Tertiary basements
0 ,o 2okra '7 ~ _"~~llllgll~lll- O,kes~s
Fig. 8. Volcanic province, dike swarms, and dominant directions of veins in the Hohi volcanic zone modified from Kamata (1989b). See
Fig. 4 for abbreviations of basins.
basin, Shishimuta and Shin-Yabakei (SY) depres-
sions (Figs. 4 and 9a). Although the buried shape
of the Pleistocene caldera (ca. 1 Ma) is the most
conspicuous feature around Shishimuta (Kamata,
1989a), the DW-6 well (Fig. 5) confirms a thick
volcanic pile under the syn-caldera pyroclastic flow,
and a Pliocene K-Ar age suggests the existence of
pre-caldera depression. The south margin of the Hohi
volcanic zone, namely the oldest trace of MTL, is
a listric surface on seismic profiles, and transforms
into a low-angle detachment beneath the volcanic
zone (Yusa et al., 1992). Thus, the left-lateral and
normal displacements on the KTL and MTL, respec-
tively, accommodated the initial N-S extension of
the study area.
Let us assume that the Hohi volcanic zone has
been buried by volcanic material compensating the
subsidence, because there is no evidence of marine
invasion since Miocene time. The burial process is expressed by the following equation:
,~ dV (1) --D = W -l ,
dt dt
where I is the lateral displacement, D is the depth of
detachment, V is the input volume of volcanic rocks,
and W is the width of the 2-dimensional model
(length of the graben). Given that the average depth
of the detachment is 3 km within the volcanic zone
(Yusa et al., 1992) and eruption rate during 5-4 Ma
is 0.38 × 10 -4 km3/year -1 per km (Kamata, 1989b),
then using the equilibrium equation the lateral ex-
tension rate (dl/dt) in the period is calculated to be
13 mm/year. This is much larger than the previous
estimate (1 mm/year; Kamata, 1989b) based on the
Pleistocene fault analysis (0.7-0.3 Ma; Muraoka and
Kamata, 1983), and suggests vigorous extension in
the initial stage (earliest Pliocene).
Y. ltoh et al./Tectonophysics 284 (1998) 135-150 147
r~
r~ ~-6
~ i ~ .=.
e~
~d
,.c
c
o
c
8
0
>~
0
148 Y. ltoh et al./Tectonophysics 284 (1998) 135-150
3.3. Migration of basins at the termination of MTL
Right-lateral motion upon the southernmost trace
0 of the MTL (Fig. 9a) deformed the ash layer hori-
zon dated as 1.34-0.3 Ma (Takemura et al., 1988)
in the uppermost part of the Sekinan Group (Yosh-
ioka, 1992). The first shift of the MTL (trace 0 to
1), therefore, occurred in early Pleistocene time or
later. After that episode, the Kuju basin in front of
the confining bend of the MTL (trace 1; Fig. 9b)
ceased to subside, while the Shonai and Beppu Bay
basins remained as active depressions (Table 1). The
Shonai Basin is filled with volcaniclastic rocks of
the Pleistocene Oita Group. This northward shift of
the westernmost part of the MTL is synchronous
with a counterclockwise shift of convergent motion
of the Philippine Sea plate (1-1.5 Ma; Nakamura et
al., 1984, 1987), which brought about an accelera-
tion of dextral slip on the MTL by highly oblique
subduction (Episode 2 in Table 1).
The second shift of the MTL (trace 1 to 2) is
probably assigned to a change of stress regime in the
Hohi volcanic zone. Volcanic rocks around the Mt.
Waita area (Fig. 9b) changed their forms of eruption
at about 0.7 Ma from lava plateaus to lava domes
accompanied by a few stratovolcanoes. Kamata et
al. (1988) explained it by local compressive stress
loaded around Mt. Waita. We assume that the change
in stress field was caused by an MTL migration
from trace 1 to trace 2 (Episode 3 in Table 1),
since then Mt. Waita has been situated in front
of the westward indenter of the fore-arc sliver of
southwest Japan. The Shonai basin ceased to subside
after this transition (Table 1) because the Yufugawa
pyroclastic-flow deposit dated to be 0.6 Ma covers
the top of Shonai basin (Hoshizumi and Kamata,
1991).
As mentioned before, the latest trace 2 of the
MTL forms the margin of active rhomboidal basin
in Beppu Bay, and contributes to the present com-
plex stress regime. At a corner of the rhomboidal
basin, namely, the intersection of the Asamigawa
and Beppu-Kita faults (see Fig. 9b), basalt activ-
ity occurred near the end of the Pleistocene epoch.
A chemical signature of the basaltic lavas obtained
from Oninomi monogenetic volcano (Ohta et al.,
1992) shows a trend which is suggestive of rift type
volcanism. Thus the strong local extension by a large releasing bend of the MTL caused not only
a depression, but also a rupture of the continental
crust accompanied by non-arc type magma activity.
It contrasts sharply with the contemporaneous is-
land-arc type eruption of the nearby Mt. Yufu and
Mt. Tsurumi (Fig. 8).
We have described the evolutionary process of
the basins within the Hohi volcanic zone. Thereby
we propose that the volcanic zone has two stages of
evolution (Table 1); the first stage of N-S extension
was triggered by the latest Miocene (6 Ma; Episode
1) resumption of the Philippine Sea plate subduction
and resultant left-lateral motions on the KTL and
parallel faults, and the second stage of migratory
rhomboidal basins was controlled by migration of
the active traces of the MTL (1.5 and 0.7 Ma;
Episodes 2 and 3) that was probably triggered by a
counterclockwise shift of the convergence direction
and oblique (dextral) subduction of the Philippine
Sea plate since early Pleistocene time.
4. Conclusions
An integrated geologic and geophysical study has
shown a series of basins forms at a bend in the ac-
tive plate margin. Latest Miocene to the Quaternary
basins in the central part of Kyushu on the boundary
between the Philippine Sea and Eurasian plates have
been generated in a volcano-tectonic depression of
the Hohi volcanic zone by the following process:
(1) Initial northward subduction of the Philip-
pine Sea plate in the latest Miocene time (~6 Ma)
activated the N-S Kokura-Tagawa Tectonic Line (K-
TL) as a left-lateral fault which bounds the western
margin of the Hohi volcanic zone (Episode 1). The
depocenter, which is now called Kuju basin, was at
the corner of the KTL and the E-W Median Tectonic
Line (MTL).
(2) Seismic reflection and borehole studies indi-
cate that the N-S crustal breakup was accommo-
dated by a low-angle detachment fault at ca. 3 km
depth throughout the Hohi volcanic zone. Assuming
that widening of the graben has been equilibrated
with deposition of volcanic rocks, a lateral extension
rate of the volcanic zone during earliest Pliocene
time of 13 mm/year is estimated to be one order of
magnitude larger than that for Pleistocene time.
(3) Relative convergence direction of the Philip-
Y ltoh et al./Tectonophysics 284 (1998) 135-150 149
pine Sea plate shifted counterclockwise in the Qua-
ternary, and right slip on the MTL was then enhanced
(Episode 2). Geologic and seismic data indicate that
the MTL has shifted twice (Episodes 2 and 3) and its
active trace migrated northward as far as 10 km in
central Kyushu. As a result, the depocenter adjacent
to the active transcurrent fault migrated northeast-
ward within the Hohi volcanic zone, namely, from
the Kuju basin (Pliocene) via the Shonai basin (early
Quaternary) to the Beppu Bay basin (late Quater-
nary).
(4) Reflecting the change of tectonic framework,
the Pliocene uniform stress with E-W aHmax in the
Hohi volcanic zone was succeeded by local tensile
and compressive regimes which were controlled by
extensional and contractional bends in the active
trace of the MTL.
Acknowledgements
We are grateful to Yasuo Takehana and Kozo
Uto for valuable discussions on regional tectonics.
Critical reviews of various stages of drafts by A.G.
Sylvester, S.E. Boyer, B.C. Burchfiel and two anony-
mous referees are appreciated.
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The Islaizd Arc (1996) 5, 289-320
Thematic Article
Anatomy and genesis of a subduction-related orogen:
A new view of geotectonic subdivision and
evolution of the Japanese Islands
YUKIO ISOZAKI
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 0-okayama,
Megur.0, Tokyo 152, Japari
Abstract The Japanese Islands represent a segment of a 450 million year old subduction-related
orogen developed along the western Pacific convergent margin. The geotectonic subdivision of
the Japanese Islands is newly revised on the basis of recent progress in the 1980s utilizing
microfossil and chronometric mapping methods for ancient accretionary complexes and their
high-P/T metamorphic equivalents. This new subdivision is based on accretion tectonics, and it
contrasts strikingly with previous schemes based on ‘geosyncline’ tectonics, continent-continent
collision-related tectonics, or terrane tectonics. Most of the geotectonic units in Japan are
composed of Late Paleozoic to Cenozoic accretionary complexes and their high-PIT metamorphic
equivalents, except for two units representing fragments of Precambrian cratons, which were
detached from mainland Asia in the Tertiary. These ancient accretionary complexes are identified
using the method of oceanic plate stratigraphy. The Japanese Islands are comprised of
12 geotectonic units, all noted in southwest Japan, five of which have along-arc equivalents in
the Ryukyus. Northeast Japan has nine of these 12 geotectonic units, and East Hokkaido
has three of these units. Recent field observations have shown that most of the primary
geotectonic boundaries are demarcated by low-angle faults, and sometimes modified by second-
ary vertical normal and/or strike-slip faults. On the basis of these new observations, the tectonic
evolution of the Japanese Islands is summarized in the following stages: (i) birth at a rifted
Yangtze continental margin at ca 750-700 Ma; (ii) tectonic inversion from passive margin to
active margin around 500 Ma; (iii) successive oceanic subduction beginning at 450 Ma and
continuing to the present time; and (iv) isolation from mainland Asia by back-arc spreading at ca
20 Ma. In addition, a continent-continent collision occurred between the Yangtze and Sino-
Korean cratons at 250 Ma during stage three. Five characteristic features of the 450 Ma
subduction-related orogen are newly recognized here: (i) step-wise (not steady-state) growth of
ancient accretionary complexes; (ii) subhorizontal piled nappe structure; (iii) tectonically
downward-younging polarity; (iv) intermittent exhumation of high-P/T metamorphosed accre-
tionary complex; and (v) microplate-induced modification. These features suggest that the
subduction-related orogenic growth in Japan resulted from highly episodic processes. The
episodic exhumation of high-P/T units and the formation of associated granitic batholith (i.e.
formation of paired metamorphic belts) occurred approximately every 100 million years, and the
timing of such orogenic culmination apparently coincides with episodic ridge subduction beneath
Asia.
Key words: accretionary complex, Japan, microfossil mapping, microplate modification, oceanic
plate stratigraphy, orogeny, paired metamorphic belts, ridge subduction, subhorizontal nappe,
Yangtze.
Accepted for publication April 1996
290 Y. Isoxaki
INTRODUCTION
The modern Japanese Islands geographically com-
prise five island arcs: the Kurile, Northeast Japan,
Izu-Bonin, Southwest Japan, and Ryukyu arcs,
showing complex patterns common in the western
Pacific. Four of these form segments of active island
arcs between the Eurasian continent and the Pacific
Ocean where the seafloor is currently subducting
westward beneath Asia; the Izu-Bonin arc is the
exception and forms an intra-oceanic arc (Fig. 1).
Ongoing subduction processes along these margins
add materials to and modify tectonic features of the
Asian continental margin. Geological studies in the
1980s revealed that the Japanese Islands evolved
under similar tectonic processes to those active
today, and that major geologic units exposed on the
islands are subduction-related products of the Late
Paleozoic to Cenozoic orogenies.
It appears that the orogenic belts in Japan have
widened oceanward by about 400 km in 450 million
years, by virtue of long-term subduction of the
Pacific seafloor along the Yangtze (South China) and
Sino-Korean (North China) continental blocks. Ac-
cording to the classic categorization of orogenic
belts by Dewey and Bird (1970), this 450 million
year old orogen of Japan corresponds to a typical
example of a Cordilleran-type orogen between a
converging pair of continental and oceanic plates,
that features a subduction complex, high-P/T
schists and coeval granitic batholiths. The Cordil-
leran-type orogen is thought to originate from
steady-state subduction of an oceanic plate beneath
a continental plate, and contrast was emphasized
between the Cordilleran-type and the Alpine-Hima-
layan style (or continent-continent collision-type)
orogen. Although great variation in orogenic styles
and possible plate tectonic mechanisms within the
so-called Cordilleran orogens has since been de-
scribed, the distinction between the Cordilleran-type
and collision-type orogens still appears justified.
Geologic and tectonic studies in Japan during the
last two decades have clarified several new and
significant aspects of oceanic subduction-related
(or classic Cordilleran-type) orogens. New observa-
tions are highlighted in the following five charac-
teristic tectonic features of Japan: (i) intermittent
growth of accretionary complexes; (ii) subhorizon-
tal nappe structure; (iii) downward younging polar-
ity; (iv) episodic formation of a tectonic sandwich
with a high-PIT unit; and (v) microplate modifica-
tion. Particularly noteworthy is the episodic (non-
steady-state) growth pattern of a subduction-
related orogen without a collision of continent or arc because it clearly contrasts with the previous
understandings on steady-state growth of the
Cordilleran-type orogen.
This article reviews the latest version of the
geotectonic subdivision of the Japanese Islands in
view of these developments and discusses its tec-
tonic implications. The implications of this review
appear to impact also on subduction-related oro-
gens in general. The new geotectonic subdivision is
fundamentally adopted from the summary by Iso-
zaki and Itaya (1991) and Isozaki and Maruyama
(1991) published in Japanese, and is slightly mod-
ified to accommodate recent information. A histor-
ical review of studies of orogeny and geotectonic
subdivision in Japan is also given in the appendix.
OCEANIC PLATE STRATIGRAPHY (OPS) ANALYSIS FOR
ANCIENT ACCRETIONARY COMPLEX (AC)
Nearly 90% of the shallow-level crust of the Japa-
nese Islands is occupied by Late Paleozoic to Meso-
zoic accretionary complexes and granitic batho-
liths. This observation suggests that the major
orogenic framework of the islands formed at this
time, and that the geotectonic subdivision of the
Japanese supracrust depends mainly on the 3D
configuration of the accretionary complexes and
their high-PIT metamorphosed equivalents.
In this article, the term accretionary complex (AC)
is strictly used in the following sense: an AC is a
geologic entity which grows in situ in trench and
trench inner wall in an active subduction zone as a
result of subduction-driven layer-parallel shortening
and vertical stacking/thickening of trench-fill mate-
rials usually composed of oceanic sediments and un-
derlying volcanic rocks. This definition excludes con-
tinental blocks or island arcs, even though they once
occurred in an oceanic domain prior to arrival at the
continental margin. The term OPS is an acronym of
‘oceanic plate stratigraphy’ and this represents a
sequence of sediments and volcanic rocks accumu-
lated primarily on an oceanic plate prior to sub-
duction-accretion at trench (Fig. 2). A full spectrum
of an ideal OPS is comprised of, in ascending order:
(i) MORB basalt; (ii) pelagic/hemipelagic sediments;
and (iii) trench-fill turbidites, similar to the rocks
recovered from drilling through modern trench
floors. In modern examples, almost identical OPS is
often preserved also in imbricated thrust packages
in the trench inner-wall, that is, the youngest part in
an AC. This unique stratigraphy was once called
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [11/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Anatomy and evolution of Japanese Islands 291
/II \ N. American Piate
Fig. 1 Plate tectonic framework around the
present day Japanese Islands Major plate in-
teractions are oceanic subductions along the
Nankai trough off southwest Japan and along
the Japan trench off northeast Japan The
former is an accretionary margin while the
latter is an erosional margin with a high con-
vergence rate and rugged-surfaced oceanic
plate The Mariana trench represents another
non-accretionary margin According to the ob-
lique subduction of the Pacific and Philippine
Sea plates southwestward moving fore-arc
slivers develop in 3 domains adding second-
ary across-arc structural features The Miocene
back-arc basin Japan Sea, is under destruc-
tion along the convergent plate boundary be-
tween Eurasian and North American plates
while a new back-arc basin is emerging in the
Ryukyus associated with rifting of the iconti-
nental crust
‘plate stratigraphy’ (Berger & Winterer 1974), how-
ever, the adjective ‘oceanic’ is added later in order to
exclude sediments derived and accumulated exclu-
sively on continental plate. Details of the concept of
OPS and the practical example of the OPS analysis
can be found in Isozaki et al. (in press b) and Mat-
suda and Isozaki (1991).
It is generally easy to recognize the occurrence of
modern AC by seismic resea,rch simply because they
occur immediately next to active trenches. Ancient
examples exposed on land, on the contrary, have
usually lost contact with their primary tectonic set-
ting nor wedge-like 3D geometry through later over-
printing processes, thus their recognition requires more specific information, not by external geometry
but by internal structure and composition. The most
reliable way to identify ancient AC, and to distin-
guish them from neighboring ones, is the ‘OPS an-
alysis’ refined mainly in Japan in the 1980s.
Ancient AC exposed on land also possess OPS
which usually consists of a sequence in ascending
order of greenstones (less than several tens of
metres thick), deep-sea pelagic chert (less than
200 m thick), hemipelagic siliceous claystone (less
than 100 m thick), and terrigenous clastic rocks
such as mudstone and sandstone (more than 200 m
thick). As most AC formed by the same teconic pro-
cess from deep-sea rocks and sediments, they are
rands1 lmudst
IC~OUS mudstone
bedded chert - birth at MOR demlae at
I s~bducllon zone t Fig. 2 Simplified ridge subduction sys-
tem and the concept of Oceanic Plate
stratigraphy (OPS) (modified from Mat-
suda & lsozaki 1991) Note the age gap
between two distinct horizons (solid tri-
angle the horizon between pillowed MORB
greenstone and pelagic chert marking the
birth of oceanic plate open triangle the
and terrigenous clastics marking the arrivdl
at trench) which represents the total travol
time of the subducting oceanic plate from
mid-oceanic ridge to trench in othu
words the age of the subducting oceanic
plate at trench auctlon zone
horizon between hemipelagic mudstone Wedge mantle
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [11/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
292 Y. Isoxaki
best distinguished by accurate age determination.
Documenting a well-dated OPS for an ancient AC
provides significant information on the timing of ac-
cretion and the age of the subducted ancient oceanic
plate. Microfossil (conodont and radiolaria) dating of
component sedimentary rocks is particularly effec-
tive in OPS analysis for ancient AC exposed on land
because it reveals (i) the accretion age at trench
( = age of the horizon between hemipelagic sediment
and trench-fill turbidite); and (ii) an age of sub-
ducted oceanic plate responsible for the accretion
( = duration of pelagic + hemipelagic deposition)
(Fig. 2). Thus if documentation of a unique OPS is
available with precise dating, it can provide a prime
identity for each ancient AC unit that appears as a
look-alike to neighboring units.
The OPS analysis for on-land exposed ancient AC
has played the main role in the drastic change in
geological studies in Japan in the 1980s (Ichikawa et
al. 1990), and the research style of detailed field
mapping in 1:5000 scale combined with microfossil
dating is here called ‘microfossil mapping’ of ancient
AC. An interval of one microfossil zone is less than 5
million years in average for the late Paleozoic to
Mesozoic, although their resolution for dating is less
precisely controlled with few tie points in time scale.
However, it is still generally quite useful to distin-
guish neighboring units in the field.
In addition, ages of metamorphosed AC of the
‘grey zone’ (see Appendix) that make microfossil
dating difficult are determined radiometrically using
various methods (K-Ar, Ar-Ar, Rb-Sr, and fission
track). In some cases, an AC is accurately dated by
two ‘radio ages’ (i.e. radiolarian-based accretion age
and radiometrically dated age of subduction-related
regional metamorphism). The subduction-related re-
gional low-grade metamorphism usually occurred
-10-20 million years later than the former (Takami
et al. 1990, 1993; Kawato et al. 1991; Isozaki in
press b), and it can add another reference feature for
comparison in OPS. This research style combining
geochronology dating and detailed field mapping is
here called ‘chronometric mapping’ of ancient AC.
Practical examples of these research styles ap-
plied to the Permian and Jurassic AC in SW Japan
are reviewed in detail in Isozaki (in press a,b) and
Kimura (1996).
GEOTECTONIC SUBDIVISION
The newly proposed geotectonic subdivision of the
Japanese Islands based on these methods (Figs 3,4)
is described here in four sections, based on domains in southwest Japan, the Ryukyus, northeast Japan,
and east Hokkaido that are at present separated
physiographically and/or tectonically by secondary
transverse faults. The fundamental scheme of this
subdivision is after Isozaki and Maruyama (1991),
and is slightly modified according to the latest infor-
mation. The Izu-Bonin Islands are not described in
this article, as they form a young island arc of intra-
oceanic nature and have lesser significance to the
primary orogenic framework along the Asian conti-
nental margin.
In southwest Japan, including Kyushu, Shikoku
and western Honshu Islands, the Cenozoic volcano-
sedimentary covers are thinner than those in other
domains due to the rapid uplift in the Quaternary,
and this allows extensive exposure of Late Paleo-
zoic to Mesozoic AC and their metamorphic equi-
valents. The Ryukyus and northeast Japan can be
essentially treated as lateral extension of south-
west Japan, however, these domains were consid-
erably modified and dislocated by secondary tecton-
ism including movement of fore-arc sliver, back-arc
spreading, and arc-arc collision. Thus these two
domains will be briefly explained after southwest
Japan as its lateral equivalents. The island of
Hokkaido is characterized by a unique setting with
an arc-arc collision between the Northeast Japan
arc and Kurile arc. East Hokkaido that belongs to
Kurile arc will be mentioned separately.
In the course of explaining geotectonic subdivi-
sion, the term ‘belt’ is used in this article to describe
the distribution of geotectonic units in two dimen-
sions. When a three-dimensional geologic entity is to
be described, non-genetic terms like unit, body,
block, complex, and nappe are used. The term ‘ter-
rane’ is avoided here because it may be confused
with a prejudicial connotation of allochthoneity (Co-
ney et al. 1980; Howell 1985), which has been sug-
gested by Sengor and Dewey (1991) and Hamilton
(1991). Refer to Ichikawa et al. (1990) for more
detailed descriptions and relevant references of pre-
Cretaceous belts; and to Taira et al. (1988,1989) for
those of much younger belts. In this long description
section, readers interested in tectonics rather than
local geology of the islands may read only the de-
scription of southwest Japan and skip those of the
Ryukyus, northeast Japan, and East Hokkaido be-
cause the latter areas represent lateral equivalents
of southwest Japan.
SOUTHWEST JAPAN
In southwest Japan the two-dimensional east-west
trending zonal arrangement of various geologic
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Anatomy and evolution of Japanese Islands 293
Fig. 3 New geotectonic subdivision of the Japanese islands (modified from lsozaki & Maruyama 1991). The geotectonic subdivision of the Ryukyus and their
correlation are shown separately in Fig. 8. Explanatory symbols for geotectonic units used in the figure and text are as follows: Southwest Japan [Rn. Renge belt ~
Sn: Sangun b ; Ak: Akiyoshi b , Mz. Maizuru b ; UT- Ultra-Tanba b (included in Mz in text), M-T. Mino-Tanba b., Ry. Ryoke b ; Sb: Sanbagawa b ; Ch: Chtchibu
b (including Northern Chichibu b.; Kurosegawa belt, and Southern Chichibu belt in text); Sh; Shimanto b. (divided into Northern and Southern Shimanto belts in
text]; Northeast Japan [Ht-Tk: Hitachi-Takanuki b. ( = Hida b.); Gs. Gosaisho b. ( = Ryoke b.), MM: Matsugataira-Motai b. ( = Renge b ); SK. Southern Kitakamt
b. ( = Oki b.); NK. Northern Kitakami-Oshima b. ( = Mino-Tanba b.); Kk: Kamuikotan b. ( = Sanbagawa b.), Hdk. Hidaka b. ( = Shimanto b.), Tokoro b
( = Sanbagawa b. + Shimanto b.); Nm: Nemuro b.]. Solid black area represents ophiolitic zone
units is more obvious if surface cover and granitic
intrusions are removed. Southwest Japan com-
prises 12 distinct geotectonic units (Fig. 5); from
oldest to youngest; a 2.0 (;a-250 Ma gneiss com-
plex; a 230 Ma intermediate-pressure type meta-
morphic complex; a 580-450 Ma ophiolite; a 400-
300 Ma high-P/T schist; a 250 Ma AC; a 230-
200 Ma high-P/T schist; a 180-140 Ma AC; a
120-100 Ma low-P/T metamorphic complex; a
100 Ma high-P/T schist, an 80 Ma AC; and a
40-20 Ma AC. These units are distributed in 15
belts, that is, from the Japa.n Sea side to the Pacific
side: Oki belt; Hida b.; 0-eyama b.; Renge b.;
Akiyoshi b.; Sangun b.; Maizuru ( + Ultra-Tanba)
b.; Mino-Tanba b.; Ryoke b.; Sanbagawa b.;
Northern Chichibu b.; Kurosegawa b.; Southern
Chichibu b.; Northern Shimanto b.; and Southern
Shimanto b. Repeated occurrence of the same unit
in more than two belts (i.e. outliers in the form of klippes and/or tectonic windows as shown in
Fig. 6) in southwest Japan causes a mismatch in
the number of belts and geotectonic units. For
example, the Jurassic AC apparently occur in three
belts, that is, the Mino-Tanba b.; N. Chichibu b.;
and Southern Chichibu b., separated from each
other for up to 50 km, although these are identical
in terms of OPS.
The three belts along the Japan Sea side, that is,
the Oki, Hida and 0-eyama belts, are intimately
linked to Precambrian crusts in nature, and they
form the core of the Phanerozoic orogen in south-
west Japan (Fig. 3). On the other hand, the other
12 belts surrounding the above three represent
zones of subduction-related accretionary growth
that account for the 450 million year-long widen-
ing and thickening of southwest Japan.
The Oki belt of continental affinity is at present
isolated from mainland Asia, as it was rifted and
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294 Y. Isoxaki
A A’ Hokkaido
B’
/ /
C’
h / /
/
S Ryukyus
D D’
\fib Pacific Ocean
C’
/ / / NE Japan
SW Japan
1 lOkm
100km
Granites a Cretaceous and
younger ACs
J Jurassic ACs
P Pre-Jurassic ACs
Continent
Shelf sediments
Fig. 4 Across-arc geologic profiles of the Japanese islands along four transects Hokkaido NE Japan SW Japan and southern Ryukyus (modified from lsozaki &
Maruyama 1991) Abbreviations for major boundary faults are as follows (I) Thrust (bold) HMT Hidaka Main Thrust HyTZ Hayachine tectonic zone Ng-Hm TL
Nagato-Hida marginal tectonic line, I-KTL Ishigaki-Kuga tectonic line BTL Butsuzo tectonic line (11) Strike-slip faults ATL Abashiri tectonic line, TTL Tanakura
tectonic line HTL Hatagawa tectonic line H-KTL Hizurne-Kesen nurna tectonic line, MTL Median tectonic line
detached by the Miocene opening of a back-arc
basin, the Japan Sea (Jolivert et al. 1994; Otofuji
1996, this issue). Judging from the lithologic/
chronologic similarity to the Precambrian rocks of
the Sobaesan massif in South Korea, the Oki belt is
regarded as an eastern extension of the Yangtze
(South China) craton (Sohma et al. 1990; Isozaki
& Maruyama 1991) which represents one of the
continental pieces rifted apart from the supercon-
tinent Rodinia at 750-700 Ma (Powell et al. 1993;
Li et al. 1995).
On the other hand, the kyanite-bearing Hida
metamorphic rocks are regarded as the northeast-
ern extension of the ultrahigh-pressure to high-
pressure metamorphic rocks along the Qinling-
Dabie suture (230 Ma continent-continent collision
zone) in central China between the Yangtze and
Sino-Korean (North China) blocks (Wang et al. 1989; Maruyama et al. 1994; Cong & Wang 1995).
The unique occurrence of Middle to Late Paleozoic
shelf strata with Boreal fauna in the periphery of
the Hida belt (Igo 1990; Kato 1990) suggests a
strong link between the Hida belt and Sino-Korean
block. The present position of the Hida belt in the
central part of Japan is due to later across-arc
contraction and juxtaposition (Komatsu 1990), and
the primary contact with the Oki belt (the Yangtze
block) has been lost. Concerning the geotectonic
correlation of the Oki and Hida belts with continen-
tal blocks and later tectonic juxtaposition, refer to
Isozaki (in press a).
The 450-580 Ma ophiolite of the 0-eyama belt
along the southern margin of the Oki belt is the
oldest oceanic material in Japan. As its easterly
extension in northeast Japan has mid-Paleozoic
sedimentary cover of continental shelf facies
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Anatomy and evolution of Japanese Islands 295
Fig. 5 Geotectonic subdivision of southwest Japan (modified from lsozaki & ltaya 1991) Hatched areas represent occurrence of Paleozoic AC, and solid black
areas indicate their tectonic outliers on the Pacific side ( = Kurosegawa b) The Nagato-Hida Marginal tectonic line (b) separates the continental Oki belt from the
younger Paleozoic-Mesozoic AC belts The lshigaki-Kuga tectonic line (f) divides Paleozoic AC belts and Mesozoic ones in southwest Japan Note its winding
surface trajectory indicating a low-angle nature as shown in Fig 6
(Ozawa 1988), the 0-eyama ophiolite probably
represents a remnant of the initial oceanic crust
that formed through the break-up of Rodinia at ca
750-700 Ma. Prior to the subduction regime
started from 450 Ma at the latest, a piece of the
initial Pacific ocean floor was likely attached to the
rifted continental margin of the Yangtze block;
that is, the Oki belt (Isoza,ki & Maruyama 1991).
The occurrence of the medium-pressure-type am-
phibolite in the 0-eyama belt probably suggests its
involvement in the 250 Ma collision event (Isozaki
in press a). The lithologic assemblage, age, and
other characteristics of these three units, with
continental affinity, are briefly described.
1, Hida belt (Hd); 1.1 Ga to 250 Ma medium-
pressure-type metamorphic rocks and 180 Ma
granites (Sohma & Kunugiza 1993). The highest
metamorphic grade reaches the upper amphibolite
facies, only locally to the granulite facies. Proto-
liths of paragneiss are composed of sedimentary
rocks that most likely accumulated along the pas-
sive continental margin. These include peralumi-
nous siliciclastic rocks and impure carbonates with
a minor amount of mafic igneous rocks. Non- to
weakly metamorphosed Middle to Late Paleozoic
shelf sequences occur fragmentally in the periphery
of the belt; 2, Oki belt (Ok): 2.0 Ga to 250 Ma
medium-pressure-type gneiss and granite complex
(Suzuki & Adachi 1994). The highest metamorphic
grade reaches the upper amphibolite to granulite
facies. Protolith is continental sedimentam rocks with minor amount of mafic igneous rocks; 3,
0-eyama belt (Oe): 450 to 580 Ma dismembered
ophiolite composed mostly of serpentinized ultra-
mafic rocks (lherzolitic.harzburgite), metagabbros,
and a fragment of medium-pressure-type amphib-
olite (Arai 1980; Kurokawa 1985).
The belts 4-15 listed below are zones of
subduction-related accretionary growth that prac-
tically account for the 450 million year-long oro-
genic widening and thickening of southwest Japan
(Figs 5,6) together with underlying granitic batho-
liths emplaced later. These AC belts including
metamorphosed equivalents were successively
added to the southeastern margin of the continent,
in particular around the Yangtze block. A summary
on OPS for these AC units in southwest Japan is
shown in Fig. 7. It is noteworthy that the timing of
accretion is generally getting younger oceanward
from the 400 Ma meta-AC to the Miocene AC, and
that the age of the subducted oceanic plate respon-
sible for accretion has been considerably variable;
from -160 million years old for the Jurassic AC to
almost zero for the Southern Shimanto AC. For a
more detailed description of belts 5-7 and 12, refer
to Isozaki (1996a); for a description of belts 8, 11
and 13 refer to Isozaki (in press b). 4, Renge belt
(Rn): 400-300 Ma high-P/T schists and associated
serpentinite (Nishimura 1990). The highest meta-
morphic grade reaches the high-pressure amphibo-
lite facies through glaucophane schist facies. The
protolith is an AC of unknown age composed of
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296 Y. Isoxnki
,- Pre-Jurassic complexes ~
of the Inner Zone 12. Ks presentj
Tectonic Outlier
Naoato- Hida
N S
1.
LA INNER ZONE OUTER ZONE -- - -.
B
Fig. 6 Geotectonic profile of southwest Japan (modified from A lsozaki & ltaya 1991, B lsozaki eta/ 1992) Profile A shows the subhorizontal nature of the piled
nappes of AC with klippes and tectonic windows on the surface Note the downward younging polarity among the AC Block diagram B is an enlargement of central
portion of profile A showing correlation of geotectonic units and boundaries across the Median Tectonic Line (M T L ) Concerning the two-fold nature of the M T L
refer to text and Figs 17b c and 18b
greenstones, siliciclastics and chert; 5, Akiyoshi
belt (Ak): Late Permian (250 Ma) AC composed of
oceanic greenstones mostly of oceanic island basalt
(OIB) origin, chert, reef limestone, and terrigenous
clastics (Kanmera et al. 1990). This unit suffered
from low grade regional metamorphism up to the
lower greenschist facies at 220 Ma; 6, Sangun belt
(Sn): 230-210 Ma high-PIT schists (Nishimura
1990). The highest metamorphic grade reaches the
high-pressure amphibolite through the glau-
cophane schist facies. The protolith is an AC,
probably containing a part of the 250 Ma Akiyoshi
AC. Neighboring schist unit with problematic 200-
180 Ma ages (probably secondarily annealed) are
also included here; 7, Maizuru belt (Mz): Middle-
Late Permian AC with 280 Ma ophiolite (Hayasaka
1990; Ishiwatari et al. 1990). The unit sometimes
discriminated as the ‘Ultra-Tanba belt’ (Caridroit
et al. 1981; Ishiga 1990) is included here. The
ophiolite suite is dismembered but its primary thickness is estimated to be -25 km; 8, Mino-
Tanba belt (MT): Jurassic AC with a minor amount
of latest Triassic and earliest Cretaceous parts
(200-140 Ma) (Wakita 1988; Nakae 1993). This
unit is composed of oceanic greenstones of OIB
origin, deep-sea pelagic chert, reef limestone, and
terrigenous clastics. Secondarily mixed AC (olis-
tostromes and melanges) occur commonly. This AC
unit is tentatively subdivided into three parts; (i)
Early Jurassic part (accreted at 200 Ma; metamor-
phosed at 170 Ma); (ii) Middle Jurassic part (ac-
creted at 170 Ma; metamorphosed at 140 Ma); and
(iii) Late Jurassic part (accreted at 140-150 Ma;
metamorphosed at 120 Ma); 9, Ryoke belt (Ry):
120-100 Ma low-PIT metamorphic rocks and asso-
ciated granites (Nakajima in press). The highest
grade part includes sillimanite-bearing gneiss. Pro-
tolith is mostly composed of Jurassic AC with
lesser amount of the pre-Jurassic AC and sedi-
ments. Age of granites ranges mostly in 120-
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Anatomy and evolution of Japanese Islands 297
UbdUCted
ceanic Plate
4- !45 Fa ra /lo n
Akiyoshi
l0w.P mel. bell
100 Ma- __ Sangun
hrgh PIT
me1 belt
0 0 ,-
0 0 N
A
A
f
-300 M: lzanagi
Mino-Tanba-Chichibu (11. 13
Iow-P mel. belt
k
400 Ma-- ~__ hanba- gawa
met. bell high-PIT
~.~ 0 e?
-0 ~ ..N a:
I,
----- 4 ._
lllllllllll
llllllllll! - Kula /Pacific /Philip.Sea
& ;I 3
Northern Shimanto ~ Southern Shimanto
low-P met bell ti
metamorphism ~ Legend
* peak Of
trench-1111 turbidites
hemipelagic sediments
pelagic sediments I
I A metagabbro I
Fig. 7 Oceanic plate stratigraphy (OPS) of AC units in southwest Japan and age of subduction-related regional metamorphism (modified from lsozaki &
Maruyama 1991) Note the oceanward and tectonically downward younging polarity not only in accretion timing (age of the horizon between hemipelagic mudstone
and terrigenous clastics) but also in high-P-T and relevant low-grade metamorphism (shown by asterisk)
70 Ma with eastward younging polarity along the
Southwest Japan arc; 10. Sanbagawa belt (Sb):
high-PIT metamorphosed Early Cretaceous AC
(Banno & Sakai 1989; Takasu & Dallmeyer 1990),
well known as the Sanbagawa schists. The highest
grade reaches the high-pressure amphibolite facies
and radiometric ages concentrate in 100-80 Ma.
The high-PIT Sanbagawa belt and the low-P/T
Ryoke belt (9) form paired metamorphic belts
(Miyashiro 1961); 11, Northern Chichibu belt (Cn):
Latest Triassic to Middle Jurassic AC equivalent to
the older part of the Jurassic AC in the Mino-
Tanba belt (8) (Hada & Kurimoto 1990). This unit
is regarded as forming a tectonic outlier of the
Jurassic complex of the Mino-Tanba belt; 12,
Kurosegawa belt (Kr): Fault-bounded mixture of
the pre-Jurassic elements (Yoshikura et al. 1990;
Isozaki et al. 1992). Components of above-
mentioned belts 3 to 7 occur chaotically as slivers,
lenses and/or blocks of various sizes and shapes,
enveloped within serpentiriite matrix. As a whole,
this unit represents a tectonic outlier of the pre-
Jurassic rocks, which occurred on the Asian conti-
nent side; 13, Southern Chichibu belt (Cs): Early
Jurassic to earliest Cretaceous AC equivalent.to the younger part of the Jurassic AC in the Mino-Tanba
belt (8) and partly to that in the Northern Chichibu
belt (11) (Matsuoka 1992); 14, Northern Shimanto
belt (Shn): Late Cretaceous scarcely metamor-
phosed AC composed mostly of terrigenous elastics
with lesser amount of oceanic rocks (Taira et al.
1988). Sporadically intervened are thin tectonic
slices of melanges that include oceanic greenstones
and bedded chert within scaly argillaceous matri-
ces; 15, Southern Shimanto belt (Shs): Paleogene
and Miocene little metamorphosed AC composed
mostly of terrigenous elastic rocks (Taira et al.
1988). A minor amount of tectonic melanges occur
in this belt.
Major geotectonic boundaries in southwest Japan
(Figs 5,6) are listed with their nature. Also noted in
parentheses are well-documented examples of these
boundary faults: (a) boundary between belts 1 and 2
(Hida b./Oki b.): low-angle thrust? (Unazuki suture)
activated probably in late Mesozoic; (b) boundary
between belts 2 and 3 (Oki b./O-eyama b.), and
between belts 2 and 4 (Oki b./Renge b.): low-angle
thrust (Nagato-Hida marginal tectonic line);
(c) boundary between belts 4 and 5 (Renge b./
Akiyoshi b.): low-angle thrust (Toyogadake thrust;
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298 Y. Isoxaki
Kabashima et al. 1994); (d) boundary between belts
5 and 6 (Akiyoshi b./Sangun b.): low-angle thrust
(Kitayama thrust), (e) boundary between belts 6 and
7 (Sangun b./Maizuru b.): unexamined; (f) boun-
dary between belts 7 and 8 (Maizuru b./Mino-Tanba
b.): low-angle thrust (Ishigaki-Kuga tectonic line);
(g) boundary between belts 8 and 9 (Mino-Tanba
b./Ryoke b.): high-angle normal fault (Arima-
Takatsuki line, Iwakuni fault); essential contact:
gradual metamorphic aureole; (h) boundary between
belts 9 and 10 (Ryoke b./Sanbagawa b.): primary
low-angle thrust (Paleo-Median Tectonic Line acti-
vated in the Tertiary) and secondary high-angle
strike-slip fault (Neo-MTL active in Quaternary). (i)
boundary between belts 10 and 11 (Sanbagawa
b./N. Chichibu b.): low-angle thrust (Sasagatani
fault; Kawato et al. 1991); 0') boundary between
belts 11 and 12 (Northern Chichibu b./Kurosegawa
b.): low-angle thrust (Agekura thrust, Nakatsu
thrust) = oceanward extension of the boundary fault
f (Ishigaki-Kuga tectonic line); (k) boundary be-
tween belts 12 and 13 (Kurosegawa b./Southern
Chichibu b.): low-angle thrust (Kanbaradani
thrust) = oceanward extension of the boundary fault
f (Ishigaki-Kuga tectonic line). (1) boundary be-
tween belts 13 and 14 (S. Chichibu b./N. Shimanto
b.): low-angle thrust (Butsuzo Tectonic Line, Tsub-
uro thrust; Sasaki & Isozaki 1992). (m) boundary
between belts 14 and 15 (N. Shimanto b./S. Shi-
manto b.): low-angle thrust (Nobeoka thrust, Aki
tectonic line).
THERYUKYUS
On the basis of strong similarity of components,
the Ryukyu Islands are basically regarded as the
southwestern extension of southwest Japan
(Fig. 8), however, a north-south trending trans-
verse fault clearly separates the Ryukyus from
Southwest Japan. This fault in west Kyushu Island
sharply cuts off the zonal arrangement of south-
west Japan. Right-lateral off-set of the Late Cre-
taceous high-P/T unit in mid- to west Kyushu
suggests the boundary fault between southwest
Japan and the Ryukyus belongs to the right-lateral
strike-slip fault system as well as the Tsushima
fault and Yangshan fault in southeast Korea (Yoon
& Chough 1995), probably formed in relation to
the Miocene opening of the Japan Sea along its
western margin.
Geological information on the Ryukyu Islands is
limited by the lack of exposures but some geotec-
tonic units are well correlated with those in south-
west Japan. Most of the units in this domain are composed of Late Paleozoic and Mesozoic AC and
their metamorphic equivalents. There is no geotec-
tonic unit with continental affinity in this domain
except Early Paleozoic ophiolite in west Kyushu. Ac-
cording to dredging data, the East China Sea is un-
derlain by Precambrian rocks that probably belong
to the Yangtze (South China) craton.
The geotectonic units hitherto known from
Ryukyus are listed below. For convenience, the
numbering and symbols that were used for units in
southwest Japan are also used to describe units in
the Ryukyus: 3, metagabbro from Nomo point:
450-580 Ma ophiolite ( = Oe) (Igi & Shibata 1979;
Nishimura 1990); 6, Tomuru metamorphic rocks:
220 Ma high-P/T schists ( = Sn) (Nishimura 1990);
8, Fusaki Formation: Jurassic AC ( = MT) (Isozaki
& Nishimura 1989); 10, Yuan Formation and
Takashima schists: Early Cretaceous AC and 60-
90 Ma high-P/T metamorphic equivalents ( = Sb)
(Ujiie & Hashimoto 1983); 14, 15, Kunchan Group:
Cretaceous and Paleogene AC ( = Sh) (Osozawa
1984).
The following two belt boundaries have been
examined on land in Ryukyus; (f) boundary be-
tween belts 6 and 8 (Sn/MT): low-angle thrust
(Ishigaki-Kuga tectonic line); and (1) boundary be-
tween belts 10 and 14 (SbIShn): high-angle fault
(Butsuzo tectonic line).
NORTHEAST JAPAN
Northeast Japan comprises northeastern Honshu
and the western Hokkaido Islands, and is sepa-
rated from southwest Japan by a left-lateral
strike-slip fault called Tanakura tectonic line
(T.T.L.), and from eastern Hokkaido by a north-
south trending fault in central Hokkaido, respec-
tively (Fig. 3). Northeast Japan is also character-
ized by an apparent zonal arrangement of several
geotectonic units on the surface (Fig. 9). This
domain has been intensely modified by secondary
tectonism, however, in particular by left-lateral
strike-slip faults relevant to the Miocene opening
of the Japan Sea. This faulting leads to uncertainty
in the primary geometry and nature of contact
among the component units. However, a compari-
son with southwest Japan can help in the recon-
struction of primary features of this domain. Com-
ponent units of northeast Japan are listed through
comparison with those in southwest Japan, and
short comments will be added for recent advances.
As well as the Hida, Oki and 0-eyama belts in
southwest Japan, the three belts in northeast
Japan, that is, the Hitachi-Takanuki b., Southern
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Anatomy mid evolution of Jupanese Islands 299
/ active trench / Ishigaki-Kuga Tectonic Line lshigaki island
(6) 220~. ~IT SSM.IS
(8) Jurassk AC
Fig. 8 Geotectonic map of the Ryukyus and their link to southwest Japan (modified from lsozaki & Nishimura 1989) Inset map (B) of East Asia shows potential
along-arc continuity of the Jurassic AC belt for example for -5000 km from west Philippines to Chind/Russia border Another 1000 km northerly extension IS
examined recently in the Kamtchatsuka Koryak region northeast Russia (refer to lsozaki in press b)
Kitakami b., and Miyamori-hayachine b., have
strong continental affinities, in particular to the
Yangtze craton and the collisional suture between
the Yangtze and Sino-Korean cratons. The rest are
AC units which later accreted to northeast Japan.
1, Hitachi-Takanuki belt (Ht-Tk): 250 Ma
medium-pressure metamorphics (Tagiri 1973; Hiroi
& Kishi 1989). The protoliths include Late Paleo-
zoic sedimentary rocks accumulated on the conti-
nental shelf and volcanic rocks of bimodal charac-
teristics ( = Hd). Granite-related thermal overprint
occurred regionally at 110 Ma which is correlated
to the Ryoke metamorphism in southwest Japan
(= Ry); 2, Southern Kitakami belt (SK): 440 Ma
granite and gneiss, 350 Ma and 250 Ma granites
( = Ok) with Middle to Late Paleozoic sedimentary covers characterized by marine fauna of Australian
(Gondwanan) affinity (Suzuki & Adachi 1993;
Kawamura et al. 1990; Kato 1990); 3, Miyamori-
Hayachine belt (MH): 450 Ma ophiolite ( = Oe) with
Paleozoic sedimentary covers (Ozawa 1988; Okami
& Ehiro 1988); 4, Matsugataira-Motai belt (MM):
300-400 Ma high-PIT schists ( = Rn) (Maekawa
1981); 8, Northern Kitakami-Oshima belt (NK-
0s): Jurassic AC ( = MT) + Early Cretaceous AC
(= Sb) (Minoura 1990; Okami & Ehiro 1988); 9,
Gosaisho belt (Gs): 110 Ma low-PIT metamorphic
rocks and granites ( = Ry). Protoliths include com-
ponents of the Jurassic AC (= MT) (Tagiri et al.
1993); 10, Sorachi-Yezo belt (SY): Early Creta-
ceous AC + 100 Ma high-PIT (Kamuikotan) schists
associated with (Horokanai) ophiolite ( = Sb); 14,
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300 Y. Isoxaki
Geotectonic subdivision
of Northeast Japan
Pacific Ocean
Mi no -Ta n b a b.)
Paleozoic shelf sediments
High-P/T schists
itachi -Takanuki b. Fig. 9 Geotectonic map of northeast Japan
(modified from lsozaki & Maruyama 1991)
Names in parentheses indicate correlative units
in southwest Japan Note that the primary sub-
horizontal geotectonic boundaries are cut by a
series of left-lateral strike-slip faults that were
activated by the back-arc basin (Japan Sea)
opening in the Miocene
Idonnappu belt (Id): Early Cretaceous to early Late
Cretaceous AC ( = Sh); 14, 15, Hidaka belt (Hdk):
Late Cretaceous to Paleogene AC ( = Sh) partly
metamorphosed into low- to medium-pressure type
metamorphics associated with 50 Ma migmatite-
granite.
Although several units are well correlated to their
counterparts in southwest Japan, some of the units in southwest Japan are apparently missing in north-
east Japan, such as Permian AC (5), 200 Ma high-
P/T schists (6), and 280 Ma ophiolite (7). Their
counterparts, however, may possibly be found also
in northwest Japan in future as subsurface under-
lying units bounded by blind thrusts, owing to the
subhorizontal structure mentioned below. The geo-
tectonic boundaries between these units in northeast
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Anatomy and evolution of Japanese Islands 301
(iii). The fault (viii) (Hidaka Main thrusts) repre-
sents another example of the secondary modification
upon the primary structure, and this transpressional
fault has been activated probably by ongoing west-
ward collision of the Kurile fore-arc sliver (Kimura
1985, 1996).
Due to severe secondary modifications, primary
orogenic structures in northeast Japan have not
been fully clarified, however, an analogy in rock
type, age and OPS of AC to southwest Japan sug-
gests that a similar subhorizontal piled nappe struc-
ture also predominates in northeast Japan which
may be represented by a series of subsurface blind
thrusts (Fig. 10). Japan are listed here: (i) boundary between belts 1
and 2 (Hitachi-Takanuki b.lSouthern Kitakami b.):
unknown; (ii) boundary between belts 2 and 3
(Southern Kitakami b./Miyamori-Hayachine b.):
vertical fault but primarily low-angle thrust (west-
ern margin of the Hayachine tectonic zone); (iii)
boundary between belts 3 and 8 (Miyamori-
Hayachine b./Northern Kitakami b.): vertical fault
but primarily low-angle thrust (eastern margin of
the Hayachine tectonic zone; Tazawa 1988).
The boundaries (ii) and (iii) appear almost vertical
in outcrop but the large-scale sinuous trajectory on
the surface suggests a potentially low-angle nature
in deeper levels (Fig. 10). The zone between these
two faults has been traditionally called the Haya-
chine tectonic zone because serpentinized ophiolite
occurs in an apparently narrow belt. The eastern
margin of this zone ( = b) corresponds to the Ishi-
gaki-Kuga tectonic line (f) in southwest Japan.
(iv) boundary between belts 1 and 9 (Hitachi-
Takanuki b./Gosaisho b.): east-dipping low-angle
fault probably activated before the intrusion of the
Cretaceous granite; (v) boundary between belts 9
and 4 (Gosaisho b./Matsugataira-Motai b.): left-
lateral strike-slip fault (Hatagawa tectonic line);
(vi) boundary between belts 4 and 8 (Matsugataira-
Motai b./Northern Kitakami-Oshima b.): left-
lateral strike-slip fault (Futaba fault); (vii) bound-
ary between belts 8 and 10 (Northern Kitakami-
Oshima b./Sorachi-Yezo b ): unknown (concealed
beneath Quaternary sediments); (viii) boundary be-
tween belts 14 and 15 (Idonappu b./Hidaka b.):
east-dipping high-angle thrust (Hidaka Main
Thrust) on the surface that translates into a low-
angle one in deeper level (Ikawa et al. 1995).
The faults (v) and (vi) are typical examples of
sinistral strike-slip faults as well as the Hizume-
Kesen’numa tectonic line that activated during the
Miocene opening event of the Japan Sea along its
eastern margin. These faults clearly cut the primary
sinuous boundary faults such as the faults (ii) and
Fig. 10 Profile of northeast Japan Note
the left-lateral strike slip faults that cut pri-
mary subhorizontal piled nappe structure
that includes inferred blind thrusts EAST HOKKAIDO
East Hokkaido has a rather complicated tectonic
history compared to the rest of the Japanese
Islands, probably reflecting its peculiar geotectonic
condition, sandwiched between two Cenozoic back-
arc basins, that is, the Japan Sea and Kurile basin
(Fig. 1). The domain boundary between northeast
Japan including West Hokkaido and East Hokkaido
is inferred in central Hokkaido in the name of the
Tokoro fault or Shibetu fault, however, its precise
position, geometry and nature are unknown owing
to thick Quaternary covers in between. Three geo-
tectonic units are recognized in East Hokkaido, as
described below (Fig. 11). There is no Paleozoic
and Early Mesozoic unit in East Hokkaido, and this
implies a unique geohistory for East Hokkaido: 13,
Yubetsu belt (Yb): Cretaceous AC; 10, 14, Tokoro
belt (Tk): Cretaceous AC and high-P/T metamor-
phic equivalents; 14, 15, Nemuro belt (Nm):
Cretaceous-Tertiary shelf sequences with unknown
basement probably composed of Mesozoic-
Cenozoic crystalline rocks of arc affinity.
These three belts are separated from each other
by north-south trending faults. The fault between
the Tokoro and Nemuro belts has a strike-slip
I
I
I
I
I Halaaawa -SW Japan---+l* NE Japan
Hayachine T Z
(3 Oe+ 4 Rni 6 Sn?)
..
10 Sb NE
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302 Y. Isoxaki
\
\ Northeast Japan -\-
ldonappu belt -3 \
\ 14,(Shimanto b.) - , i \
younging polarity --+
Northern Kitakami-
Oshima b.
8. (Mino-Tanba b.) __ East Hokkaido ~
North
10 Sorac hi-Y ezo b.
.(Sanbagawa b.)
~ Yubetsu b. (Shimanto b.) 14.
\'V \
\,Hidaka b?.
\\ (shimanto b.) 100 km
14, 15.
Fig. 11 Geotectonic map of Hokkaido (modified from Kimura unpubl data, 1996) Note the polarity reversal of younging direction in AC units across the Hidaka
Main Thrust which represents a collision suture between the Kurile arc to the northeast Japan arc
nature, and this fault, called the Abashiri tectonic
line, has activated in a right-lateral manner during
the opening event of the Kurile back-arc basin.
Refer to Kimura (1996) for further details.
The main problem in Hokkaido lies in how to
interpret the origin of the parallel-running two
coeval blueschists belts, that is, the Kamuikotan b.
of northeast Japan and the Tokoro belt of East
Hokkaido. The opposite younging polarity in AC
between northeast Japan and East Hokkaido
(Fig. 11) suggests that a secondary collision/
amalgamation has doubled the primarily linear
blueschists belt along the Asian margin. Several
other interpretations were also proposed but a final
agreement has not yet been reached. GEOTECTONIC HISTORY OF THE JAPANESE ISLANDS
The time-space relationships among the orogenic
units described here suggest that the Japanese
Islands have grown oceanward by almost 400 km
across the arc since the mid-Paleozoic. On the basis
of the new geological data and the revised sub-
division of the Japanese Islands, the geotectonic
history of the Japanese Islands is summarized
below. The history began with rifting of a super-
continent in the late Neoproterozoic, and was fol-
lowed by a tectonic inversion shifting from an
extensional (Atlantic-type) regime to a convergent
(Pacific-type) regime around 500 Ma. Since then an
oceanic subduction-related accretion regime has
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Anatomy and evolution of Japanese Islands 303
the core of Japan, which began as a segment of the
Yangtze continental margin. In addition, a frag-
ment of the Sino-Korean continental block also
participated in the orogenic growth of Japan after
the 250 Ma collisional event (Isozaki in press a). persisted, and is responsible for the oceanward
growth of the islands. Figure 12 summarizes this
history from the birth of the islands at ca 700 Ma
to their separation from Asia in the Miocene.
BIRTHPLACE OF PROTO-JAPAN These two continental blocks, particularly the
Yangtze block, have played the most important roles
The Oki belt in southwest Japan and the Southern
Kitakami belt in northeast Japan represents part of in constraining the configuration of the Japanese
Phanerozoic orogen.
I _- -
rifling ,-
I 1.0-07Ga I Asthenospheric mantle
li I1
Break-up of supercontinent Rodinia I and birth of Pacific (750-700 Ma)
(b)
Y~JW~' Passive marain sediments Remnant of orimilive
Onset of initial oceanic subduction
(tectonic inversion) (-500 Ma)
Calc-alkaline 0-eyama fore-arc ophtolite
volcanic arc --, ,, Omanward accretionay growth
in the Paleozoic (400-250 Ma)
lCollision of YangtzeISino-Korea
Hida gneiss
(collision comDlex) Ok Oe Rn Ak 6n.c: Mz V.T Sb SIB
system along continental
Fig. 12 Series of simplified cartoons (a-g) showing 700 million years of geotectonic evolution of the Japanese Islands (a) The birth of Japan along a Proterozoic
rifted continental margin of Yangtze by the break-up of the supercontinent Rodinia and the birth of the Pacific Ocean at 750-700 Ma. (b) After widening of the
proto-Pacific Ocean basin, tectonic inversion occurred to initiate an intraoceanic subduction zone off the rifted Yangtze margin around 500 Ma, leaving a small fragment
of primitiveoceaniccrust (0-eyamaophiolite). (c) Thearc-trench system wasestablished by ca450 Maand gaveacalc-alkalineoverprint on thefore-arcophiolite. (d)
The subduction zone matured to accommodate the oldest AC which corresponds to the protoliths of the 400 Ma high-P/T Renge schists. The high-P/T Renge belt
together with coeval granite belt (fragmented) form the oldest set of paired metamorphic belts. (e) Successive oceanic subduction widened the accretionary edifice
during the Late Paleozoic (the Renge and Akiyoshi belts). On the opposite side of the Yangtze block, another continental block (Sino-Korea) collided and thrust over the
Yangtze at around 250 Ma. (f) Oceanward growth on the Pacific side continued during the Mesozolc and Cenozolc to widen and thicken the accretionary edifice in the
form of subhorizontal piled nappes. The accretionary growth was punctuated by the episodic arrival of mid-oceanic ridge roughly every 100 milion years, which
generated 3 to 4 sets of sandwich structure of high-P/T nappe between low-pressure AC units above and below (e.9. Sangun and Sanbagawa belts; see text for details).
(9) By the back-arc opening of the Japan Sea at 20 Ma, Japan detached from mainland Asia; a continental arc becamean island arc The Hida belt represents a remnant
of the 250 Ma collision suture between the Yangtze and Sino-Korean blocks that partly features an ultrahigh-pressure metamorphic belt.
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304 Y. Isoxuki
The relation of the Yangtze and Sino-Korean
continental margins to protoJapan goes back to
nearly 750-700 Ma when it is believed that the
supercontinent Rodinia started rifting apart (Hoff-
man 1991; Dalziel 1992; Powell et al. 1993). Judg-
ing from the patterns of radial dyke swarms and
rifted basins (Park et al. 1995; Bond et al. 1984),
the breakup of Rodinia was probably triggered by a
superplume rising to the surface from the mantle/
core boundary. Consequently, the proto-Pacific
ocean was born in 750-700 Ma by the rifting of
Rodinia, and several continental fragments, includ-
ing the Sino-Korean and Yangtze blocks were
dispersed in various directions (Fig. 12a). Due to
similarities in Neoproterozoic to Cambrian strati-
graphy, Laurentia (North America) and Australia
are probably the conjugate continental block(s) to
the rifted Yangtze margins (Li et al. 1995). How-
ever, the paleoposition of the Sino-Korean block
with respect to Rodinia remains highly enigmatic.
The birthplace of Japan was probably located some-
where in the northern periphery of Rodinia in mid-
latitudes at about 700 Ma. Most of the dispersed
continental fragments once again assembled to
form another (semi-)supercontinent, Gondwana-
land, about 500 Ma (Hoffman 1991; Dalziel 1992),
however, some large blocks such as Laurentia, Sibe-
ria and Baltica were isolated from the superconti-
nental mass. Likewise, the absence of evidence for
late Neoproterozoic to Cambrian collision in China
suggests that both the Sino-Korean and Yangtze
blocks, including proto-Japan, were also isolated
from Gondwanaland, although Early to mid-Paleo-
zoic faunal provincialism suggests that the Yangtze
block (including proto-Japan) and Australia were
close neighbors (Burrett et al. 1990; Kato 1990).
Due to limited exposure and later tectonic modi-
fication in Japan, there is little evidence to conclude
that continental rifting, such as extensional fault
system, rift-related bimodal volcanism, and rift-
related sedimentary sequences were active at this
time. Stratigraphical and paleontological studies
suggest that the small distribution of Early to
Middle Paleozoic (Ordovician to Devonian) terrige-
nous clastic/carbonate sequences in the peri-
phery of the Hida belt, Hitachi-Takanuki belt,
Southern KitakamiIMatsugataira-Motai belt, and
the Kurosegawa belt all represent remnants of
continental shelf facies accumulated along the
rifted Proterozoic continental margins. The Korean
peninsula and/or mainland China may have pre-
served such features of Rodinian rifting, however,
further structural and stratigraphical analyses are
needed. INVERSION FROM PASSIVE TO ACTIVE MARGIN
With the exception of Precambrian gneissic clasts
in younger sediments, the oldest unit of oceanic
affinity in Japan is the 580 Ma ophiolite in the
O-eyama belt, southwest Japan. Its occurrence in
the periphery of the Oki belt (ancient Yangtze
margin) suggests that this unit is a remnant of
proto-Pacific oceanic crust. The O-eyama ophiolite
and its equivalent in the Miyamori-Hayachine belt
in northeast Japan have a bimodal distribution of
radiometric ages; one at 580 Ma and another at
480-450 Ma (Ozawa 1988; Nishimura & Shibata
1989). Isozaki and Maruyama (1991) explained the
age distribution of the oldest ophiolite in Japan in
the following way.
The tectonic history of the O-eyama ophiolite is
two-fold: (i) following rifting with the emplacement
of nascent oceanic crust, a MORB-like oceanic
crust formed at the proto-Pacific mid-oceanic ridge
around 580 Ma (Fig. 12a); and (ii) these rocks
were then intruded around 450 Ma by calc-alkaline
volcanism of arc affinity when a new intra-oceanic
subduction was initiated (Fig. 12b,c).
O-eyama is regarded as the only example of a
fore-arc ophiolite in Japan, while other ophiolitic
rocks are regarded as accreted fragments of an-
cient seamounts, rises and plateaus (Isozaki et al.
1990b; Kimura & Maruyama in press). Between
580 Ma and 480 Ma, tectonics in proto-Japan
changed dramatically from riftlridge-related ex-
tension to subduction-related compression, and
passive margins changed rapidly to active margins.
This tectonic inversion probably corresponds to
global plate reorganization, in particular to the
opening of the proto-Atlantic (Iapetus) ocean on
the opposite side of the globe. Accretion of the
O-eyama ophiolite can be explained by either of the
following two mechanisms: (i) initiation of a land-
ward dipping subduction zone within the primary
oceanic crust (see Fig. 12b,c); or (ii) collision of an
island arc system from the ocean side and a
reversal of subduction polarity.
ACCRETIONARY GROWTH OF THE JAPANESE ISLANDS
After the tectonic inversion around 500 Ma, proto-
Japan began a state of accretionary growth that
persists today. Within 50 million years after initi-
ation of intra-oceanic subduction, the arc-trench
system featured AC, high-PIT schists (metamor-
phosed AC), and granitic batholiths (Fig. 12d,e)
that is, in the Renge belt plus Kurosegawa belt
(=tectonic outlier of the former) in southwest
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Amtomy and evolution of Japanese Islands 305
Sino-Korean and Yangtze blocks collided along the
Qinling-Dabie suture, generating an ultrahigh-
pressure metamorphic belt, best preserved in the
Dabie mountains and Shandon province in China
(Wang et al. 1989; Maruyama et al. 1994; Cong &
Wang 1995). The region of ultrahigh-pressure
metamorphism appears to be restricted to the
central part of the 3000 km long suture, suggest-
ing that a promontory collision may have locally
generated ultrahigh-pressure conditions (Isozaki in
press a). East of the Shandon peninsula, the coeval
250 Ma regional metamorphism is detected solely
in the Hida and Hitachi-Takanuki belts in Japan;
they are characterized by medium-pressure-type
metamorphism. Although these belts lack ultra-
high-pressure metamorphism, from a geographical
viewpoint (Fig. 13), they may also correspond to an
eastern extension of a collision-related metamor-
phic belt (Sohma et al. 1990), including the Imjin-
gang or Ogchon zones in Korea (Ernst et al. 1988;
Lan et al. 1995). The Hida belt and its equivalent
occur at the eastern terminal of the suture. Data
from this belt indicate a decrease in grade and
magnitude of a collision-related metamorphic belt.
After 250 Ma, accretionary growth in Japan radi-
ated from a core composed of the amalgamated
Sino-Korea plus Yangtze blocks. Thus while minor
in extent, the Japanese Islands do contain elements
of a non-accretionary, continent-continent collision-
type orogen. Japan, and in Southern Kitakami and Matsuga-
taira-Motai belts in northeast Japan. In particular,
the oldest AC in Japan is the protolith of the 450-
400 Ma high-P/T schists. The high-P/T schists and
coeval granitic rocks are elements of a paired
metamorphic belt (with a high-P/T belt on the
ocean side and a low-P/T belt on the continent
side; Miyashiro 1961).
Following the oldest 450 Ma unit, Late Paleo-
zoic, Mesozoic and Cenozoic AC were formed
through subsequent subduction. At least several
major oceanic plates have subducted beneath the
Yangtze margin, leaving more than 10 distinct AC
belts. Numerous oceanic fragments derived from
subducted oceanic plates, including deep-sea sedi-
ments and seamount-derived basaltslreef lime-
stone, were accreted to Japan. Details of the
accretion processes during the Permian and Juras-
sic periods are reported in Isozaki (in press a,b).
Accretionary growth apparently has been not
continuous. Including the youngest AC now under
construction along the Nankai trough off southwest
Japan, total accretionary growth is nealy 400 km
in across-arc width (Fig. 12f), not taking into
account the material loss by subduction-erosion.
Thus the overall AC-dominated orogen in Japan
has grown oceanward for almost 400 km during
the 450 million years (-100 km per 100 million
years).
As all of the AC units in Japan were formed in
situ by subduction along the Yangtze (South
China) continental margin, they are autochthonous
to Asia, with the exception of small oceanic frag-
ments peeled off from subducting oceanic crust and
accreted landward into AC (Isozaki et al. 1990b).
It thus appeared that the Japanese islands do not
represent a collage of ‘suspect or exotic terranes’
that had existed prior to subduction and accretion
processes (Coney et al. 1980).
It was a mere 20 million years ago when the
Japanese Islands obtained their present configura-
tion through back-arc spreading (Fig. 12g). How-
ever, the accretionary growth of the Japanese
Islands will likely continue until other continental
blocks, such as Australia or North America, collide
against Asia to form a future supercontinent
(Maruyama 1994).
REMNANT OF CONTINENT-CONTINENT COLLISION
While most of the geotectonic units in Japan are
the results of oceanic subduction, the Hida belt and
the Hitachi-Takanuki belt are remnants of
continent-continent collision. At about 250 Ma, the DISCUSSION
The newly revised geotectonic subdivision and re-
constructed tectonic history of the Japanese Is-
lands clarifies the properties of the Cordilleran-type
orogenic growth. Five tectonic features are de-
scribed following these observations in Japan. They
are: (i) step-wise growth of AC; (ii) subhorizontal
nappe structure; (iii) downward younging polarity;
(iv) a tectonic sandwich of high-P/T metamorphic
units; and (v) secondary tectonic modifications.
These may represent the principal characteristics
of subduction-related orogenic belts in general.
STEP-WISE GROWTH OF AC UNITS
The 450 million-year-old subduction-related history
of the Japanese Islands is characterized by the
intermittent formation of AC, clearly detected by
OPS analysis (Fig. 7). Several intervals lack AC
(Carboniferous-Early Permian, Early-Middle Tri-
assic, and late Early Cretaceous), and this empha-
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306 Y. Isoxaki
Jurassic AC
sizes the overall zonal (piled nappe) arrangement
of AC in Japan (Figs 5,6,8-11). Secular changes in
relative plate motion between the Asian continent
(or its precursory fragment) and subducted oceanic
plates may be responsible for such episodicity
because AC cannot form from a highly oblique
subduction or in an along-arc strike-slip regime
(Maruyama & Sen0 1986).
On the other hand, extensive modern AC can be
constructed where a large amount of sediments are
supplied and the subduction rate is moderately low.
No accretion or tectonic erosion (subduction-
erosion) occurs when a trench is starved of sedi-
ments or the subduction rate is too high (von
Huene & Lallemand 1990; von Huene & Scholle Fig. 13 Tectonic framework of 250 Ma
East Asia showing a continent-continent col-
lision orogen between the two Precambrian
cratons that is the Sino-Korean and
Yangtze blocks (modified from lsozaki &
Maruyama 1991) The collision suture fea-
turing the 250 Ma ultrahigh-pressure rneta-
morphic belt of the Qinling-Dabie zone in
China extends eastward to the Hida belt and
Hitachi-Takanuki belt in Japan character
ized by the 250 Ma medium-pressure meta-
morphism The pathway of the suture in the
Korean peninsula is still controversial as two
alternative interpretations are proposed, one
along the Ogchon zone and the other along
the Imjingan zone
1991). Thus the apparent episodicity in the forma-
tion of AC in Japan may also indicate the episodic
preservation of AC rather than formation. None-
theless such episodic growth of AC is a primary
characteristic of oceanic subduction-related oro-
gens that persist for 100 million years or longer.
The formation of high-PIT metamorphic and
coeval granitic belts also appears episodic as these
occur in highly restricted time intervals. As they
occur roughly every 100 million years, their forma-
tion may also be episodic and/or periodic.
In contrast, the classic concept of the Cordil-
leran-type orogeny proposed by Dewey and Bird
(1970) assumed a steady-state orogenic process,
including the simultaneous formation of subduction
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Anatomy and evolution of Japanese Islands 307
Subhorizontal piled nappe structures generally
characterize continent-continent collision-type oro-
gens, like the European Alps or the foreland fold-
and-thrust belt in the Appalachian and Canadian
Rocky mountains. Development of similar struc-
tures in an arc-trench setting dominated by sub-
duction tectonics is noteworthy and should be
tested in other AC-dominated orogens older than
100 million years. complexes ( = AC), granitic batholiths, regional
metamorphic belts, and relevant geologic struc-
tures. Later publications, h.owever, emphasized the
non-steady-state nature of the arc-subduction zone
setting (Dewey 1980). Although various interpre-
tations are given, causes for the episodicity in
subduction-related orogeny has not been well clar-
ified. A possible cause for the episodic exhumation
of high-PIT metamorphosed AC and coeval gra-
nitic activity associated with low-P/T regional
metamorphism will be discussed later.
SUBHORIZONTAL PILED NAPPE STRUCTURE OF AC
Most of the geotectonic boundaries between adja-
cent AC units in southwest Japan are low-angle
faults, and therefore the AC units are considered to
be piled nappes (Fig. 6). Recent microfossil/
chronometric mapping documented the occurrences
of regional klippes and windows of these nappes
even in thickly vegetated areas. These observations
indicate that the subhorizontal piled nappe struc-
ture governs the fundamental tectonic framework
of the 450 million-year-old orogen in southwest
Japan (this study: see also Hara et al. 1977;
Charvet et al. 1985; Faure 1985). Primary oro-
genic structures in the Ryukyus and northeast
Japan have not yet been fully mapped but a
subhorizontal piled nappe structure controlled by
subsurface blind thrusts has been predicted
(Tazawa 1988; Isozaki & Nishimura 1989;
Fig. 10).
To a first approximation, such a subhorizontal
piled nappe structure is consistent with subduction-
related deformations, particularly to the subhori-
zontal shortening caused by underplating through
which new materials are added to the sole of the
previously formed accretionary wedge through
step-wise activation of a subhorizontal decolle-
ment. The size of an individual ancient AC nappe is
nearly 200 km in width across the arc, similar to
the size of the widest modern AC wedge (von
Huene & Scholle 1991).
The documentation of a predominant subhorizon-
tal structure in Japan is contrary to the traditional
view that vertical tectonics dominated horizontal
tectonics. There are, in fact, some along-arc verti-
cal faults of strike-slip nature (Neo-M.T.L.,
T.T.L.), but most of them became active in the
Cenozoic and were driven by microplate activities.
The total amount of displacement along the verti-
cal fault, however, is too small to account for the
present zonal arrangement of belts in Japan that
extend along the arc for more than 1000 km. DOWNWARD YOUNGING POLARITY IN AC NAPPES
The along-arc zonal arrangement of these ancient
AC is most clearly demonstrated in southwest
Japan (Fig. 14a). Remarkably, these AC show an
oceanward younging polarity in map view without
windows and klippes. In addition, ages of regional
metamorphic belts and granitic batholiths also sug-
gest an oceanward younging polarity. This appar-
ent polarity is a function of the piled nappe struc-
ture of the AC (Figs 6,12f). Downward younging
polarity is also recognized within an individual AC
nappe; this was clearly demonstrated in the Juras-
sic AC, which comprises six or more subnappes
(Isozaki in press b). Such oceanward and down-
ward younging polarity is consistent with the
growth patterns of modern AC.
Although detailed information is scarce, the
Ryukyus and northeast Japan plus East Hokkaido
also appear to display younging polarity. The
present zonal arrangement in northeast Japan,
however, does not fit with the oceanward younging
polarity in southwest Japan (Fig. 14b). This is due
to secondary strike-slip faulting related to the
opening of the Japan Sea in the Miocene which
modified the primary orogenic configuration.
TECTONIC SANDWICH OF HIGH-P/T AC UNIT
There are three geotectonic units composed of
high-P/T schist in the Japanese Islands formed at
around 450-300 Ma, 200 Ma, and 100 Ma. The
450-300 schists can be further subdivided into two
distinct units. All of these high-P/T units also
occur as subhorizontal nappes (Fig. 6). The most
striking feature of the high-P/T nappes is their
sandwich-like structure, where a high-PIT nappe is
tectonically interleaved between two unmetamor-
phosed AC nappes (Maruyama 1990; Isozaki &
Maruyama 1991). The nappes composed of high-
P/T metamorphosed units can be traced laterally
for more than 500 km along the arc, even though
they are usually thinner than 2 km. For example,
the 100-70 Ma Sanbagawa schists (Sb) occur as a
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308 Y. Isoxaki
SW-Japan / Ryukyus NE Japan
Sino - Korean crafon
Precambrian craton collision zone
ophiolite / high-P/T metamorphosed AC non- to weakly metamorphosed AC Shn
\
\
low-PrT metamorphosed part
Fig. 14 Cartoon showing stepwise oceanward growth of southwest Japan with a high-P/T sandwich structure after removing windows and klippes (see text) in
a map view and a contrasting feature in northeast Japan which was secondarily modified by strike-slip shredding lo obscure the primary structure
nappe sandwiched between the overlying Jurassic
AC (MT) and the underlying Cretaceous AC (Sh)
(Kawato et al. 1992; Sasaki & Isozaki 1992).
There are high pressure gaps between the high-
P/T nappe in the middle of the sandwich and the
adjacent AC nappes. These gaps, which are greater
than several kilobars in pressure, correspond to
10-20 km of crustal thickness. In order to preserve
these pressure gaps without metamorphic anneal-
ing, the tectonic insertion of the high-P/T nappe
must have been rapid and caused by tectonic
exhumation of a high-PIT nappe into a low-
pressure domain. In fact, all of the high-P/T
nappes are bounded by a set of subhorizontal faults
along their top and bottom surfaces, however, the
sense of dislocation between these faults is oppo-
site. To compensate for the pressure gap between
the hanging wall and foot wall, the upper fault is
normal while the bottom fault is reverse. Synchro-
nous activation of such paired faults is believed to
result in the insertion of the thin high-P/T nappe
into the low-pressure domain.
These observations strongly suggest that the
exhumation of the high-PIT nappes was tectonic and episodic rather than caused by steady-state
processes wherein buoyant uplift of a metamorphic
domain together with over- and underlying un-
metamorphosed units was proposed. This episodic
tectonic exhumation mechanism called ‘wedge ex-
trusion’ was first suggested by Maruyama (1990;
see also Maruyama in press). A very similar model
was also proposed for Himalayan medium-pressure
gneisses by Burchfiel et al. (1992). The downward
younging polarity among AC is not disturbed by
the intermittent intercalation of high-P/T nappes.
This suggests that subduction-driven burial of the
protolith AC, high-PIT metamorphism, and tec-
tonic exhumation of metamorphosed AC all oc-
curred in the same structural horizon, likely along
the Wadati-Benioff plane. The large-scale sandwich
(containing a thin well-done steak) with downward
younging polarity cannot be formed through any
other known exhumation process (Suppe 1974;
Cloos 1982; Platt 1986).
Based on radiometric ages determined by various
methods, the timing of metamorphic peak temper-
ature and the subsequent cooling history of high-
P/T units in Japan (Itaya & Takasugi 1989; Nishi-
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Anatomy and eTolution qf Japanese Islands 309
gests an eastward along-arc younging polarity for
these intrusions (Nakajima et al. 1990; Kinoshita
1995). This observation is consistent with the
oblique subduction of the Kula/Pacific ridge be-
neath Asia as mentioned above (i.e. northeastward
passage of the TTR triple junction of the Eurasia/
Kula/Pacific plates, because the subducted ridge-
related slab window may have been a heat source).
The coincident timing of high-P/T nappe exhuma-
tion and granite belt formation suggests a causal
relationship to ridge subduction. Further research is
needed to elucidate the mechanism of the subhori-
zontal ejection of high-PIT nappes under a buoyant
subduction regime induced by the movement of
young oceanic crust. The preservation of high-PIT
units in the fore-arc is possible even under high heat
flow from the subducted ridge if the exhumation of
the high-PIT nappe preceded the arrival of the
ridge-crest at the trench. The metamorphic/cooling
ages of the high-P/T unit (100-70 Ma) are slightly
older than the subduction timing of the brand new
oceanic plate (ca 75 Ma), and this may support the
interpretation provided above.
Prior to the 100-70 Ma event, exhumation of the
high-PIT nappe occurred probably three times in
the 600 million year history of the Japanese Islands
(Fig. 15). Such episodic occurrences of the high-
PIT sandwich structure suggests that these were
also consequences of episodic ridge-subduction in
the Paleozoic and early Mesozoic (Fig. 16). The
distribution of pre-Cretaceous granites in Japan is
highly limited; most of them are fragmented.
Nonetheless the ages of the granites that are
preserved suggest episodic formation of a granitic
belt in the intervals of 450-400 Ma, 350 Ma, and
250 Ma.
If those three pre-Cretaceous geologic episodes
are regarded as the result of episodic ridge subduc-
tion, the tectonic history of the Japanese Islands mura et al. 1989; Shibata & Nishimura 1989;
Takasu & Dallmeyer 1990), the exhumation of
high-P/T nappes occurred episodically. It is note-
worthy that these episodic events apparently coin-
cide with the formation of granitic batholith in
low-P/T metamorphic belts on the continent side
(Isozaki & Maruyama 1991; Fig. 15). The best
example of such paired metamorphic belts, de-
scribed by Miyashiro (1961), is the 100-70 Ma
high-P/T Sanbagawa belt on the trench side and
the low-P/T Ryoke belt on the arc side of south-
west Japan.
The formation of the Cretaceous paired belts was
probably related to subduction of the Kula/Pacific
ridge along the Asian margin. An observation on
contemporary OPS in mid-Cretaceous AC in the
Shimanto belt in southwest Japan (Taira et al.
1988) suggested a progressive decrease in the age
of the subducted slab in the Late Cretaceous
(Fig. 16). At about 75 Ma, an AC was formed by
subduction of a brand new oceanic plate (a mid-
oceanic ridge per se).
In addition, based on hot spot tracks, the recon-
structed paleo-plate motion in the Pacific domain
over the last 150 million ,years indicates that the
Kula-Pacific mid-oceanic ridge subducted obliquely
beneath the eastern Asian continental margin
around 100-70 Ma (Engebretson et al. 1985).
Thus a TTR (trench-trench-ridge) triple junction
may have migrated from southwest to northeast
off-shore of Japan.
The phenomenon of ridge-subduction and its
tectonic effects on the continental margin have long
been discussed (Pitman & Hays 1968; Wilson
1973; Uyeda & Miyashiro 1974); however, its
tectono-thermal effects and orogenic consequences
are still controversial (Farrar & Dixon 1993; Thor-
kelson 1994). A recent compilation on the age of
Mesozoic granite emplacement in East Asia sug-
Stepwi$e oceanward propaRation of BS belt and granite belt Fig. 15 Simplified diagram of geotectonic
profile of the Japanese Islands (modified
from lsozaki & Maruyama 1991) Note the
oceanward and tectonically downward N
younging growth of AC nappes incluijing
high-P/T blueschist (BS) nappe Formation
of the sandwich structure of BS nappe occurs
intermittently roughly every 100 million
years, accompanying coeval granite batho-
lith belt with low-P/T metamorphic belt
which is usually positioned -100-200 km
continentward from coeval BS belt The iriter-
mittent oceanward propagation of the BS belt
and granite belt appears to have occurred
when mid-oceanic ridges episodically col- b
High-PIT metamorphic AC (BS) nappe
Low-pressure AC iiappe
Granite intrusron /
tided and subducted beneath proto-Japan
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310 Y. Isoxaki
Age (Ma)
Formation of
high-PIT sandwich
(tectonic exhumation)
Age of subducted 400 300 200 100 0 ,
I
450-40018 330-280Ma 230-210~a 100-70Ya
Renge I1 Sang"" Sanbsgawa 'engel * * * *
-~ I
I/ --0 1
,' (\ I, io
15 slabdeduced from
OPS of AC
interacted oceanic plate
Ridge subduction
can be simplified to 450 million years of oceanward
accretionary growth punctuated four times by
ridge subduction. For each event, granites and
associated low-P/T metamorphic belts formed on
the continent side of Japan, probably separated by
200 km or more from the coeval high-PIT schist
belt (Fig. 15). The nearly 100 million year period-
icity in sandwich formation may correspond to the
consumption of a major oceanic plate dissected by
mid-oceanic ridges on both sides. The Japanese
Islands experienced subduction of at least five
major oceanic plates (from young to old, Pacific,
Kula-Izanagi, Farallon, an unnamed older oceanic
plate, and the plate which originated during initial
rifting; Fig. 16).
The anatomy of the subduction-related orogen in
Japan and the tectonic interpretation of these
structures strongly contradict earlier views of
'Cordilleran-type' orogeny. The most significant
difference lies in the recognition of episodicity in
orogenic activity related to ridge subduction. Pre-
vious models for exhumation of high-P/T units
assumed a steady-state and long-term process re-
lated to subduction of the normal ocean floor. For
such an oceanic subduction-related orogenic pro-
cess involving episodic culmination by ridge sub-
duction, Isozaki and Maruyama (1 99 1) introduced
a new term, 'Miyashiro-type orogeny'. The name is
given after Akiho Miyashiro's outstanding contri-
butions to subduction zone tectonics, particularly
the first perception of paired metamorphic belts
(Miyashiro 1961) prior to the birth of plate tecton-
ics in the mid-l960s, and a keen perspective on
ridge subduction and relevant geologic phenomena
in the 1970s (Uyeda & Miyashiro 1974). ,' '\\ 7
- LOO ,' ?
~ 150
-- 200Ma-
we7 Unnamed Farallon $1 Izanagi ,,*" Kula $1 Pacificp
?nn 77n
SECONDARY MODIFICATION: MICROPLATE TECTONICS
The primary orogenic structures of the Japanese
Islands are controlled by the spatial arrangement Fig. 16 Timetable of ridge subduction and
orogenic culmination in Japan indicated by
formation of the high-P/T sandwich structure,
age of subducted oceanic slab deduced from
OPS analysis and paleoplate interaction
based on hot spot track analysis (modified
from Isozaki & Maruyama 1991) The names
for oceanic plates that interacted with Japan
are adopted from the reconstructed plate in-
teraction in the Pacific domain by Engebret-
son eta/ (1985) Note the temporary coinci-
dence among the zero age of subducted
oceanic plate at 75 Ma exhumation of the
high-P/T Sanbagawa metamorphic belts, and
passage of KulaiPacific ridge off Japan
of main components (i.e. ancient AC, regional
metamorphic rocks and granitic batholiths).
Subduction-related orogens in Japan are character-
ized by subhorizontal piled nappes of AC with
downward younging polarity, and by the episodic
occurrence of a sandwich structure of high-PIT
nappes and unmetamorphosed AC. The primary
orogenic features, however, were modified or de-
stroyed in some cases by secondary tectonic pro-
cesses. In particular, microplate-related tectonics
has the profound ability to secondarily reorganize
primary structures (Miyashiro 1982). In Japan,
there are three important secondary tectonic pro-
cesses (Fig. 17; fore-arc sliver movement, back-arc
basin opening, and arc-arc collision) that are all
orogenic manifestations of microplate activities
(Isozaki 1989; Isozaki & Maruyama 1991).
A fore-arc sliver is a decoupled microplate of a
frontal arc driven by a strike-slip component of the
oblique subduction of an oceanic plate. In the
Japanese Islands, there are three active fore-arc
slivers (the South Ryukyu sliver, Nankai sliver in
southwest Japan, and Kurile sliver in East Hok-
kaido; Fig. 1). Along-arc movement of fore-arc
slivers can create three distinct features that de-
stroy the primary orogenic edifice (i.e. along-arc
strike-slip fault on are-side margin, and across-arc
compressional structure in front, plus an exten-
sional one in rear of the sliver). The best example
of an along-arc strike-slip fault is the Quaternary
Neo-M.T.L. in the Kii peninsula and on Shikoku
Island that demonstrates a remarkably linear sur-
face trajectory for more than 500 km (Figs 3,5).
This right-lateral strike-slip fault was driven by
the westward movement of the Nankai fore-arc
sliver (Fig. 17b), cutting the older low-angle fea-
ture of the Tertiary paleo-M.T.L. (Isozaki 1989;
Yamakita et al. 1995; Fig. 6). This westward sliver
translation is also associated with across-arc com-
pression along the Bungo Strait between Shikoku
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Anatomy and evolution of Japanese Islands 311
(a) Arc-arc collisian (b) Fore-arc sliver
Philippine Sea Plate
(c) Opening of back-arc basin I’ A\
Fig. 17 Three representative modes of secondary modification of primary orogenic structure related to microplate tectonics (modified from lsozaki & Maruyama
1991). (a) Formation of syntaxis in the older accretionary orogenic system and accretion of exotic arc crust occur at an arc-arc collision front around the Izu
peninsula where the Izu arc penetrates almost perpendicularly to the Southwest Japan arc (b) Formation of along-arc strike-slip fault (Neo-M T.L ) and across-arc
extensional and compressional structure by along-arc movement of the Nanakai fore-arc sliver (c) Formation of along-arc extensional structure (Miocene rifted
basins with bimodal volcanism), compressional structures (thrusting along paleo-M.T L ) and the strike-slip fault (T T.L.) occurred when the Japan segment
detached from mainland Asia through the rifting-opening of the Japan Sea. Note these Neogene to Quaternary structures have modified considerably the pre-existing
major structures of the ca 450 Ma accretionary orogen
and Kyushu Islands, and across-arc extension in
Ise Bay. Across-arc compression related to a fore-
arc sliver is best observed in the elevated Hidaka
mountains in central HoIkkaido. Kimura (1985)
explained the exposure of lower crustal rocks of
the Hidaka belt (Komatsu et al. 1989) as a tectonic
manifestation of westward frontal collision of the
Kurile fore-arc sliver to northeast Japan. Seismic
reflection research (Ikawa et al. 1995) recently
documented the crust-cutting detachment surface
that dips eastward from the Hidaka main thrust
(Fig. 11).
Back-arc spreading is another tectonic process
that has occurred frequently in the western Pacific
since the latest Mesozoic, and is also a powerful
modifier of primary orogenic structures in Japan.
For example, when the Japan Sea opened in the
Miocene, it split pre-existing orogenic structures
into several blocks (Figs 9,10,14). Although a cer- tain amount of rotation was involved, the opening
of the Japan Sea basin has been attributed to
dislocation of a pair of north-south running strike-
slip faults along the basin margin (Fig. 17c). The
chaotic alignment of belts in northeast Japan is
primarily due to the series of left-lateral strike-
slips along the eastern margin of the Japan Sea.
The T.T.L. presently dividing southwest Japan and
northeast Japan, and the parallel Futaba and Hata-
gawa faults are typical examples of an eastern
margin strike-slip fault that obliquely dissects an
older along-arc zonal arrangement. In contrast, a
segment of the right-lateral strike-slip fault mark-
ing the western margin in east Korea was recently
noted (Yoon & Chough 1995). In addition, a do-
main of compression in the fore-arc may have been
associated with back-arc spreading (Fig. 17c). It is
difficult to explain the side-by-side juxtaposition of
the low-P/T Ryoke metamorphic belt and high-PIT
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312 Y. Isoxaki
Sanbagawa metamorphic belt in modern southwest
Japan (Fig. 5). These belts should have been sepa-
rated from each other by at least 100-200 km
when they formed in the Late Cretaceous arc-
trench system (Fig. 18a). The origin of the low-
angle Paleo-M.T.L. between these two belts
(Fig. 6) may be related to development of a hori-
zontal detachment surface in the arc crust and to
trenchward dislocation of the upper crust along this surface, probably by back-arc spreading
(Isozaki & Maruyama 1991; Fig. 18b). For further
details on the opening of the Japan Sea, refer to
Jolivert et al. (1994), Otofuji (1996) and Yamash-
ita et al. (1996). The present-day opening of the
Okinawa trough (Kimura et al. 1988) is still in a
nascent stage of back-arc spreading (Fig. 1). This
extensional regime appears to propagate north-
ward into mid-Kyushu (note the Beppu-Shimabara
6 c usually 100-200 km (100-701 a? ca 50 km ___b.
Akryorhr-Sangun A B C
pre-Juraralc
S posslble
external llmlt of
volc. tront trench
- -
high-P/T Sanbagawa met.
Mlno-lanba
Juraralc complex
100 km
subducted
Kula-Paclflc
rldge
Orlglnal conflguratlon of the Cretaceous
palred metamorphlc belta in SW Japan p i~
- n'B' = Intra-arc shortenlng 1 = dlslocatlon 01 delslchmenl Iault I 130-40 Ma] fore-arc contraction 1
A', B' C'
100 Ma trench 100 Ma voic front
trench Paleo-MTL volcanlc [rant
Shlmanto
Palaogana complex back -arc
extension
mlsalnp volume of the Ryokc fore-arc rcplon
= accreted volume 01 the Shlmanto complex
Secondary juxtapositlon of
hlgh-P/T and low-P/T motamorphlc belts
Fig. 18 Schematic cross-section showing the primary configuration of the Cretaceous arc-trench system in southwest Japan (a) and secondary modification
involving intra-arc shortening associated with back-arc (Japan Sea) spreading (b) Note the juxtaposition of the Cretaceous paired metamorphic belts induced by
the trenchward translation of upper arc-crust along an intracrustal horizontal detachment surface
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Anatomy and evolution of Japanese Islands 313
tica, and eastern Australia, as the Paleo-Pacific
ocean floor was subducted along these continental
margins after the 750-700 Ma breakup of Rodinia.
Reconnaissance studies (Sedlock & Isozaki 1990;
Isozaki et al. 1992; Isozaki & Blake 1994) includ-
ing a compilation of previous works for the
Klamath-Franciscan belt in California (Isozaki &
Maruyama 1992; Maruyama et al. 1992) suggests
fundamental similarities between the geotectonic
evolution of California and that of Japan. It ap-
pears promising to extrapolate the above-described
new aspects and analyzing schemes for ancient AC
to other orogenic belts of different time-space
co-ordinates in the Earth’s history, including the
Precambrian. graben with active Aso and Unzen volcanoes),
accordingly, Kyushu Island, now composed of sev-
eral geotectonic units, may eventually split into
two.
Arc-arc collisions are the third process that can
modify the primary geotectonic structures in Ja-
pan. The ongoing collision of the Izu-Bonin arc
against the southwest Japan arc is a typical exam-
ple of this process (Fig. 17a). The northward buoy-
ant subduction of an intra-oceanic arc can indent
into a pre-existing orogenic structure, leaving a
clear V-shaped mark called ‘orogenic syntaxis’, in
the west of Tokyo (Fig. 4). In addition, the accre-
tion of arc crust basement has been achieved in a
step-wise manner, adding a significant volume in
central Japan (Taira et al. 1989). Several tectonic
blocks of intra-oceanic arc origin around Mt. Fuji
(Amano 1986) thus represent bona fide ‘allochtho-
nous or exotic terranes’ in Japan.
CONCLUSION
The Japanese Islands represent a segment of a
Phanerozoic subduction-related orogen developed
along the western margin of the Pacific Ocean.
Significantly, this orogen has grown oceanward
nearly 400 km during the last 450 million years.
The anatomy of the Japanese orogen, best pre-
served in southwest Japan, is characterized by a
subhorizontal piled nappe structure, which involves
multiple AC nappes including those of high-P/T
metamorphosed AC. Orogenic growth is unusual in
that it includes: (i) step-wise growth of AC units;
(ii) tectonically downward younging polarity; and
(iii) intermittent sandwich structure of high-PIT
nappes. Episodic subduction of mid-oceanic ridges
( =migration of TTR triple junctions) appears to
explain the episodic exhumation of high-P/T
nappes and the formation of granite/low-P/T
metamorphic belts every 100 million years. Micro-
plate tectonic processes such as the movement of
fore-arc slivers, back-arc spreading and arc-arc
collisions, secondarily modified the primary struc-
ture of Japan.
In order to establish a general model for
subduction-related orogenic processes, the new
geotectonic model for the origin of Japan needs to
be tested on other orogens formed in similar tec-
tonic environments. For example, the occurrence of
a subduction-related orogen characterized by a
similar anatomy would be expected in Circum-
Pacific Phanerozoic orogens in western North
America, western South America, Western Antarc- ACKNOWLEDGEMENTS
The author would like to thank S. Maruyama, Y.
Nishimura, T. Itaya, T. Matsuda and many stu-
dents from Yamaguchi University for a decade-
long discussion in various tectonic aspects of the
Japanese Islands. Sincere thanks are also due to
A. J. Kaufman, S. Maruyama, G. Kimura, and
L. Ivany who constructively reviewed the manu-
script. P. F. Hoffman gave valuable comments on
the neo Proterozoic distribution of continental
blocks.
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as refractory residue in the Sangun-Yamaguchi zone,
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or virtue. In Dewey J. F., Gass I. G., Curry G. B., Harris
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Hida nappe and the tectonics of Mesozoic sediments:
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318 Y. Isoxaki
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APPENDIX: HISTORICAL REVIEW ON STUDIES OF
OROGENY AND GEOTECTONIC SUBDIVISION OF THE
JAPANESE ISLANDS
In order to establish a geotectonic subdivision for a
particular region, conventional field mapping on a
regional scale is essential, and this usually requires
painstaking work by numerous geologists over an
extended period of time, particularly in thickly
vegetated regions like the Japanese Islands that
are at a humid mid-latitude. In the 100-year his-
tory of geological research in Japan, many efforts
have been made to synthesize all available geolog-
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and these all came together with the revised sum-
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pp. 88 (in Japanese).
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Jurassic radiolarian assemblages from the Inuyama
area central Japan. Journal of Geosciences Osaka
City University 23, 135-55.
YOON S. H. & CHOUGH S. K. 1995. Regional strike slip
in the eastern continental margin of Korea and its
tectonic implications for the evolution of Ulleung
Basin, East Sea (Sea of Japan). Gdogical Society qf
America Bulletin 107, 83-97.
YOSHIKURA S., HADA S. & ISOZAKI Y. 1990. Kurosegawa
Terrane. In Ichikawa K., Mizutani S., Hara I., Hada S.
& Yao A. eds. Pre-Cretaceous Terranes of Japan,
pp. 185-201, Publication of IGCP #224, Osaka.
Japanese Islands can be roughly divided into four
stages, based on the orogenic concept that gov-
erned the understanding of contemporary geolo-
gists; these are: (i) the pre-geosyncline concept
stage (1865-1940); (ii) the stage of importing the
geosyncline concept (1941-1955); (5) the stage
of popularization of geosyncline-based orogeny
(1956-1975); and (iv) the stage of plate tectonics-
based orogeny (1976-present).
The first stage is represented by activities of E.
Naumann, B. Lyman and other foreign geologists
who imported modern geology as well as other
sciences to Japan immediately after the Meiji revo-
lution in 1868, after two century-long diplomatic
isolation. During this stage, foreign and domestic
geologists recognized fundamental structures of
the Islands, particularly major tectonic boundaries
such as the Median Tectonic Line (M.T.L.) and
Itoigawa-Shizuoka Tectonic Line (I.-S. T.L.); how-
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [11/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Anatomy and evolution of Japanese Islands 319
the importance of distinguishing orogenic phase in
H. Stille’s sense. His summary was the first appli-
cation of the then world-popular geosyncline con-
cept to a terra incognita named Japan. His
achievement, however, marks a prime milestone in
geological studies in Japan and his subdivision
provided a foundation for later works. One point to
note is his emphasis on nappe structures through-
out southwest Japan, as this idea was revived in
the 1990s in the plate tectonic framework.
The third stage was governed by geologists of
the post-World War I1 generation who strongly
criticized Kobayashi’s summary. They emphasized
the higher resolution and accuracy of their data set
through detailed mapping and the advantage of
their new tectonic models (Yamashita 1957: Minato
et al. 1965: Ichikawa et al. 1970), however, their
perspective was more or less the same as that of
Kobayashi, that is, popularization of the classic
geosynclinal concept with minor modification in the
context of regional geology. Although the world
according to fixism was still there, the quality of
the regional geological description of the Japanese
Islands improved much during this stage, with the
use of megafossil dating, particularly fusulinid
biostratigraphy. Regional distribution of geotec-
tonic units and locations of mutual boundaries were
clarified, and the geotectonic subdivision made in
this stage appears very similar to the present one
in map view except for ‘vertical ’ major boundary
faults (Fig. 19b). Accordingly, sporadically exposed
older mid-Paleozoic rocks, including high-grade
metamorphic rocks and serpentinites, were all ex-
plained as geosynclinal basement rocks squeezed
out from deeper levels along putative ‘crust-
penetrating vertical faults’. In the early 1970s,
several avant-garde geologists from Japan were
the first to propose plate tectonic interpretations
for the evolution of the Japanese Islands, with
special emphasis on subduction-related tectonics
(Matsuda & Uyeda 1971: Uyeda & Miyashiro
1974). The majority of scientists, however, still
held conservative understandings linked to the geo-
syncline concept, and the geotectonic subdivision
was not revised from the plate tectonic viewpoint.
The most significant reform in the geotectonic
subdivision occurred in the fourth stage. Nearly a
decade after the fundamental construction of plate
tectonics in the late 1960s, younger Japanese
geologists started to prefer mobilism rather than
fixism. Down a long and winding road with many
arguments, including the conversion from a geo-
syncline world to a plate tectonics world, both in
individuals and in society (Kanmera 1976), re- ever, the overall geotectonic subdivision of the
Islands was still a rough sketch. In hindsight, this
era up to the 1930s in Japan may be called a time
of fundamental ‘find-and-describe’ in preparation
for the following stage of importing the concept of
the geosyncline.
During the second stage, T. Kobayashi was the
first Japanese geologist to synthesize a grand view
of the geotectonic evolutim of the Islands on the
basis of the stratigraphic and megafossil age data
(Kobayashi 1941). He recognized most of the im-
portant geotectonic units currently known, but his
subdivision (Fig. 19a) was strongly biased by the
geosyncline concept. In particular, he emphasized
Akiyoshi Orogen Sakawa Orogen
~~~~~~
MTL
IctoScalel Opa Nappa
:b) lchikawa era/ 1970
N Shimanto Orogen S ~- Honrhu Orogcn ___
Cp Mz MT Ry Sb Ch Hdk Mr
km MTL
[i 30 . ,50 km, ,
Fig. 19 Classic examples of geotectonic map and profile of the Japanese
Islands (modified from (a): Kobayashi 1941, (b): lchikawa et a/. 1970).
Compare these classic geotectonic subdivisions and profiles backgrounded by
geosynclinal viewpoint with the current version shown in Figs 3,4 Before
1975, major geotectonic units in Japaii were all explained as Precambrian
sialic basement and overlying Paleozoic to Mesozoic geosynclinal sediments.
Hd. Hida b , Ok: Oki b., Cg: Chugoku b., Mz, Maizuru b., MT: Mino-Tanba b.;
Ry: Ryoke b.; Sb: Sanbagawa b.; Ch: Chichibu b.; Hdk. Hidakagawa b. ( = Sh:
Shimanto b.); Mr. Muro b.( = NK. Nakamura b.); Km: Kitakami marginal b.;
KI: Kitakami b ~ Sm. Soma b.; Ab: Abukuma b., As: Ashio b.; Jo: Joetsu b.
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [11/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
320 Y. Isoxaki
gional geologic information was constantly accu-
mulated on a nationwide basis. The great reform
came in two waves, the first one in the early 1980s
and the second in the early 1990s.
The first wave was brought by two factors com-
bined: (i) the remarkable enhancement in micro-
fossil (conodont and radiolaria) biostratigraphy, in
particular the high resolution dating of the pre-
Cretaceous or ‘so-called Paleozoic eugeosynclinal
sedimentary rocks’ in Japan (Koike 1979; Isozaki
& Matsuda 1980; Tanaka 1980; Yao et al. 1980,
Yamato-Omine Research Group 198 1; Nakaseko et
al. 1982); and (ii) the acceptance of a subduction-
accretion concept (Kanmera 1980; Taira et al.
1983) which was constructed mainly through deep-
sea drilling and regional seismic profiling on mod-
ern AC. As a result, most of the ‘so-called Paleo-
zoic eugeosynclinal rocks’ in Japan were revealed
to be Jurassic AC by the mid-1980s. Age of
accretion for each AC unit was precisely dated, and
this enabled mutual discrimination of neighboring
and similar-looking Paleozoic to Cenozoic AC in
Japan. This reform initiated a considerable redraft-
ing of the geotectonic history of the Islands in
terms of accretion tectonics, as well as their geo-
tectonic subdivision. Utility of microfossil dating
for ancient AC and the relevant results in geotec-
tonic subdivision of Japan up to the late 1980s are
summarized in Ichikawa et al. (1990).
On the other hand, concerning tectonic inter-
pretation, the ‘allochthonous or suspect terrane’
concept (Jones et al. 1977; Coney et al. 1980;
Howell 1985) invaded Japan in the early 1980s
and allowed many geologists in Japan to believe
the occurrence of exotic continental and/or oceanic
blocks and to use the term ‘terrane’ to describe
various orogenic units (Saito & Hashimoto 1982;
Mizutani 1987; Ichikawa et al. 1990). On the
geotectonic subdivision, the ‘terrane’-based under-
standing of the attitude of boundary faults
between the orogenic units is of note because most
of the ‘terrane boundaries’ were regarded as
vertical faults of a strike-slip nature related to
‘terrane dispersion’ (Taira et al. 1983). The nappe
tectonics, on the contrary, was revived also in
the early 198Os, and its significance with subhori-
zontal boundary faults was emphasized by French
and domestic geologists (Hara et al. 1977;
Yamato-Omine Research Group 1981; Faure 1985;
Charvet et al. 1985; Hayasaka 1987). French
geologists, in particular, interpreted the nappe-
related subhorizontal structure as a result of an
ancient continent-microcontinent collision, how-
ever, evidence for the putative collided micro- continent per se was not persuasive. Independent
from these works, relative plate motions were
partly reconstructed for the late Mesozoic and
Cenozoic Pacific region by Engebretson et al.
(1985), and the correlation between the plate
interactions in East Asia and orogenic events re-
corded in Japan was first discussed by Maruyama
and Sen0 (1986). Owing to the mixed effects of
these various interpretations and existence of prob-
lematic ‘gray zones’ mentioned below, understand-
ing of the 3D structure and tectonic evolution of
the Japanese orogen was in a state of confusion in
the 1980s.
The second wave in the early 1990s that pro-
vided the final tool for redrawing the subdivision of
the Islands came in the form of success in chrono-
metric dating of weakly metamorphosed AC. AC
well-recrystallized by regional metamorphism, such
as the Sanbagawa blueschists, had already been
dated by radiometric methods in the 1970s, how-
ever, less recrystallized metamorphic AC had not
been dated at all owing to difficulty in mineral
separation techniques. Similarly, while microfossil
analysis is powerful in dating non- to weakly
metamorphosed AC, those metamorphosed to the
greenschists facies were mostly left untouched
owing to difficulty in microfossil extraction. Such a
dilemma had left a considerable number of undated
weakly metamorphosed AC units, the ‘gray zones’,
in Japan even after the 198Os, and this delayed the
completion of the geotectonic subdivision of the
Japanese Islands. In the late 1980s, an advanced
technique of fine-grained mineral separation was
developed (Nishimura et al. 1989) that solved the
‘gray zone’ problem (Isozaki et al. 1990a, 1992;
Isozaki & Itaya 1991; Suzuki et al. 1990; Takami
et al. 1990; Kawato et al. 1991). Not only did this
permit the dating of the greenschist facies AC
rocks, but it allowed subdivision of high-PIT re-
gional metamorphosed units on the same basis as
non- to weakly metamorphosed ones (Isozaki &
Maruyama 1991). By 1993, most of the important
geotectonic boundaries were clearly defined by age
difference and were re-examined in the field. This
clarified that these boundary faults, that is, the
nappes, are essentially subhorizontal, except for
several vertical ones of secondary origin. In other
words, there is no deep crust-cutting vertical fault
zones nor a collage of exotic terranes in Japan. The
predominance in subhorizontal structures of the
Japanese Islands are currently accepted by many,
and are utilized, for example, in the latest version
of a geologic map of the islands issued by the
Geological Survey of Japan (1992).
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Isozaki (1996) Anatomy and genesis of a subduction_related orogen A new view of geotectonic.txt
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Earth and Planetary Science Letters, 92 (1989) 335-346 335
Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands
[21
Nd and Sr isotope systematics of Miocene to Holocene volcanic rocks
from Southwest Japan: volcanism since the opening of the Japan Sea
Paul A. Morris 1 and Hiroo Kagami 2
i Department of Geology and Geophysics, University of Sydney, N.S.W. 2006 (Australia)
2 Institute for Study of the Earth's Interior, Okayama University, Misasa 682-02, Tottori-ken (Japan)
Received August 17, 1987; revised version accepted January 30, 1989
Isotope (Nd and St) and geochemical data for twelve volcanic rocks from the Shimane-Old (San'in) region of
Southwest Japan give some insight into the Miocene-Recent volcanic history of southern Japan since rifting from the
Asian mainland. The oldest rocks examined (Dolerites and Ushildri Formation) were erupted during the opening of the
Japan Sea at about 15 Ma, and comprise arc tholeiites derived from a depleted source (end > 0; %, < 0). 11 Ma old
basalts of the nearby Matsue Formation have transitional arc-within-plate trace element patterns, and end < 0 and
%r > 0, interpreted in terms of a depleted mantle source modified by a slab-derived fluid. Volcanics of Old Dozen
Island in the Japan Sea are alkalic, derived from an enriched mantle plume, isotopically distinct from nearby Old Dogo
Island. Eruption of the calc-alkaline Wakurayama Andesite (early Pliocene) and both arc and within-plate-type basalts
( < 1 Ma) show that magmas were tapped from a variety of sources. Taking into account plate tectonics and seismic
data, the subduction influence until the Quaternary was the Pacific plate, with Philippines Sea plate subduction
affecting Quaternary and younger volcanies.
1. Introduction
We present Nd and Sr isotope data and whole-
rock geochemistry for volcanic and hypabyssal
rocks outcropping over a limited area of southwest
Honshu and on nearby Oki Dozen Island in the
Japan Sea. As only twelve isotopic analyses were
made, we stress that this is a reconnaissance study
but maintain that the availability of K/Ar dates
and whole-rock chemistry for each sample pro-
vides valuable data for interpreting the Miocene
to Recent volcanic history of Southwest Japan.
Initial investigations of igneous rocks in the
Shimane-Oki region (e.g. [1,2]) have identified
Miocene to Recent calc-alkaline, tholeiitic and
alkalic basalts occurring in close proximity. These
rocks range in age from 15 Ma to less than 0.5 Ma
(i.e. Miocene to Holocene) and although rocks of
diverse composition occur adjacent to each other,
they often have the same eruptive age [3]. Thus,
the Shimane-Oki area is a fertile ground for in-
vestigating isotopic variations in rocks of diverse
composition outcropping in a restricted area that
have been erupted over a short time interval.
0012-821x/89/$03.50 © 1989 Elsevier Science Publishers B.V. Furthermore, we will attempt to integrate these
variations into the tectonic development of South-
west Japan from the Miocene onwards.
Lastly, the twelve isotope analyses presented
here supplement the rather meagre Nd-Sr isotope
data base currently available for Japanese volcanic
rocks [4-9].
2. Tectonic framework
In order to present a coherent framework in
which to examine chemical and isotopic varia-
tions, we present a short overview of the Miocene
to Recent plate tectonic history of Southwest
Japan,
The Japanese islands are currently underlain by
two subducting plates (Fig. 1), the Pacific plate
(PP) and the Philippines Sea plate (PSP). The
volcanic front for the PSP is outlined by volcanic
activity on Kyushu (e.g. Aso Volcano) and South-
west Honshu (e.g. Mts. Sambe and Daisen). Until
recently this was the only evidence for the extent
of the PSP beneath southwest Honshu as there
was no deep focus seismic activity in the south-
336
A
.~. "z::.~ ~..,. .,'
°', ' i pp I ..;' ".' #
• J,.JI : L=, Damen ,~) '~ .~.. I o_ Shimane-Ok . L,~ ~" I~
Sambe --~"
~"~'x~" \\ ',%
~" ...~,~ \\ ',p,
\ oo
I/,o~i~ ......... - \%
/ /~- / ~ Shikoku \'~ / /_.o / ~ = .... \o a,o~ / ~ ,,/o,:.v u a~ln \'7
/ ~9 /, x,,- PBP \ " ~N." / .'~ .... \
.,'.~ o soo B !-~ OKI DOGO
.:.,,..I
SEk
43/;!:":: ..... 47. ~e~2 ~
64p~ pEblIt~SU LA ~"
Matsu6~** 1 ~~-'~, .:..::".:. 8 .:"" (.."2 '..~b...
9 L s
km
e37
Fig. 1. A. Location of Shimane-Oki region in Southwest Japan and present-day plate tectonic configuration of Japan [10]. Trenches:
short broken lines; volcanic fronts: solid lines. Depths to Benioff zone (long broken line) shown in km. PSP = Philippines Sea plate•
PP = Pacific plate. B. Sample locations. 1, 64 = Dolerites; 43, 47 = Ushildri Formation; 17, 18 = Matsue Formation; 50, 51, 56 = Oki
Dozen Volcanics; 24 =Wakurayama Andesite; 37=Yasugi Basalt; 15 = Dalkon Jima Basalt. More precise locations and petro-
graphic descriptions are given elsewhere [27].
west Honshu region. However, detailed seismic
network data [10] showed that the PSP leading
edge was currently beneath the Shimane region:
the lack of seismic data for larger earthquakes was
interpreted in terms of the PSP being "seismically
dead", the plate having broken off, or subduction
only being intermittent.
PSP subduction began approximately 40 Ma
ago [11], with subsequent inter-arc spreading (terminating at 16-17 Ma [12,13]) resulting in
formation of the Shikoku Basin [14,15]. As the
plate boundary southwest of Honshu during the
Miocene was a transform rather than a subduction
margin [11] (Fig. 1), it is unlikely that PSP sub-
duction had any influence on volcanic activity in
the Shimane region until the Quaternary. Thus,
any Miocene subduction-related volcanism in the
Shimane area is related to the PP subduction.
Moreover, shallow earthquake focii (ScSp) data
indicate that the PSP has not reached the Old
Islands, so any subduction influence in the Japan
Sea is also PP related. Although sketchy, seismic
observations have identified an inclined surface
approximately 30 km thick at a (shallow) depth of
50-70 km beneath the Shimane district, interpret-
ed as the upper surface of the downgoing PSP
[10,161.
Recently published work on the paleomagne-
tism of late Cenozoic rocks of southwest Japan
have confirmed the Middle Miocene as a crucial
period in the evolution of the Japanese islands
[17-20] indicating that the Japan Sea opened by
about 57 ° clockwise rotation from the Asian
mainland at approximately 15 Ma. Opening was
completed in less than 1 Ma [21]. Paleomagnetic
data from the Omori Formation (overlying the
Ushikiri Formation) shows the same declination
as younger rocks, confirming this short period of
back-arc spreading (T. Watanabe, personal com-
munication). Several lines of evidence suggest that
the Japan Sea is, at least in part, ensialic: con-
tinental crust has been identified in the eastern
Japan Sea [22]; 2 Ga gneisses have been recorded
from Southwest Japan [23,24], and gneissic gravels
have been recorded from Shimane Peninsula [25].
Opening of the Japan Sea occurred during the
effusive volcanism of the Middle Miocene (" Green
Tuff Movement" [26]) and the oldest rocks ex-
amined in this paper coincide with the separation
of Japan from the Asian mainland. In the follow-
ing, we consider the chemistry of Miocene-Recent
volcanics in a tectonic framework.
3. Sample selection
Samples for isotope analysis were selected from
36 rocks analysed for major and trace elements
([27] and unpublished data). In order to allow the
extra flexibility of investigating isotope variations
with age, we ignored samples for which K/Ar
dates were not available. Detailed petrographic
descriptions and precise locations are given
elsewhere [10] and only a summary is presented
below and in Fig. 1.
All samples (apart from volcanic rocks from
Old Dozen Island and basalt from Yokota) out-
crop in close proximity (20 x 18 km) on Shimane
Peninsula and near to Matsue City in Southwest 337
Honshu (Fig. 1). The oldest rocks (Dolerites and
Ushikiri Formation ranging in age from 12.8 to
14.6 Ma [3]; samples 1, 64, 43, 47) are submarine
eruptives and hypabyssal rocks occurring along
the length of Shimane Peninsula, ranging from
dolerites, gabbros and basalts (grouped as
Dolerites by Kobayashi et al. [1]) to basalts and
rhyolites of the Ushikiri Formation. They com-
prise varying proportions of clinopyroxene,
plagioclase and opaques, with orthopyroxene in
rhyolites. In keeping with their eruption in the
"Green Tuff" region [26,28] they contain variable
amounts of secondary chlorite, carbonate and
sericite. However, samples analysed in this study
contain less than approximately 3 vol.% secondary
minerals [27].
Isolated outcrops of compact, massive bluish-
black basalt outcrop in and close to Matsue City,
comprising the Matsue Formation [29]. Extensive
examination of scattered outcrops in and near
Matsue City indicate that there are two petro-
graphic and geochemical groups [27,29] that have
the same age range (10.7-11.1 Ma [3]). The first
group (represented by sample 17 in this study)
carries hornblende phenocrysts and groundmass
biotite, whereas the second (sample 18) has minor
olivine, more abundant clinopyroxene, and minor
groundmass biotite. Of the two islands comprising
the Oki Group (60 km north of Shimane
Peninsula), the larger one (Old Dogo) has been
extensively examined in terms of its petrology
[8,30-32], as there is a diverse suite of upper
mantle and lower crustal xenoliths. In terms of
petrography, and major and trace element chem-
istry, alkaline volcanics from Old Dozen are simi-
lar to the alkalic basalts and trachyte differenti-
ates of Old Dogo Island [27,33] although there are
differences in isotope chemistry. K/Ar dating of
Old Dozen volcanics and intrusives [3] reveals that
alkaline volcanism began earlier (at 6 Ma) on Old
Dozen, the later stages overlapping with the al-
kaline volcanic activity on Old Dogo [34].
The 5 Ma Wakurayama Andesite is the only
calc-alkaline rock examined in this study, forming
a prominent outcrop near Matsue City. It com-
prises rare hornblende phenocrysts set in a flux-
ioned groundmass of feldspar laths with rare bio-
tite flakes.
Isolated outcrops of Holocene (1.2 Ma) basalts
occur south and east of Matsue City near Yasugi
338
and Yokota (e.g. 37). These basalts (termed the
Yasugi Basalt here) are grey massive to columnar
jointed rocks comprising hornblende and biotite
microphenocrysts in a felsic groundmass. One
sample from a quarry near Yasugi contains dis-
aggregated lherzolite xenoliths.
The youngest samples examined are from
Daikon Jima Island (Daikon Jima Basalt; sample
15) in Nakaumi Lagoon east of Matsue City.
These occasionally vesicular basalts comprise
clinopyroxene plus plagioclase with minor
opaques. K/Ar age is less than 0.5 Ma.
4. Analytical procedures
Major and trace elements were measured by
XRF at Shimane and Sydney Universities respec- tively. Reproducibility was checked using several
international reference standards and found to
give good agreement between observed and ex-
pected values. Analytical procedures for Sr and
Nd isotopes at the Institute for Study of the
Earth's Interior are detailed in Appendix 1 and
Table 2.
5. Geochemistry
5.1. Major element chemistry (Table 1)
The three Oki Dozen samples (56, 51, 50) and
the two Matsue Formation basalts (17, 18) are
clearly alkaline in terms of alkalies/silica (Fig.
2A), whereas Dolerites (1, 64), Ushikiri Formation
samples (43, 47) and the Wakurayama Andesite
(24) are subalkaline. The two youngest rocks (15:
TABLE 1
XRF major element oxide and trace element analyses for samples analysed for isotopes
1 64 43 47 17 18 50 51 56 24 15 37
SiO 2 52.51 45.78 52.91 71.21 52.77 48.17 65.23 52.73 49.94 63.21 48.97 50.48
TiO 2 1.47 0.73 1.26 0.90 1.52 1.96 0.36 1.81 2.66 0.35 1.94 0.85
A1203 16.57 16.87 18.23 13.66 16.46 17.46 16.97 19.07 18.12 17.49 15.59 16.92
Fe203 5.38 4.37 4.07 1.25 4.10 4.24 2.48 3.41 3.98 1.90 4.07 3.09
FeO 4.78 4.71 4.57 1.55 2.81 3.65 0.95 3.07 4.51 2.05 6.11 4.09
MnO 0.29 0.20 0.17 0.06 0.09 0.13 0.08 0.13 0.14 0.07 0.16 0.13
MgO 3.19 8.10 3.55 0.14 3.70 5.52 0.17 2.88 3.26 1.78 6.87 9.49
CaO 7.50 11.06 8.87 3.03 7.58 8.05 0.97 6.30 7.72 5.20 9.45 8.89
Na20 3.42 2.03 3.58 4.32 4.09 3.61 5.54 4.19 3.80 4.12 3.07 3.61
K20 0.73 0.31 0.85 1.78 3.36 2.06 5.95 3.28 2.48 1.17 1.07 1.03
P20~ 0.22 0.10 0.26 0.31 1.20 0.67 0.04 0.69 0.76 0.13 0.30 0.32
LOI 3.03 4.83 2.24 1.40 2.17 4.38 2.04 2.56 2.38 1.57 1.45 0.94
Total 99.09 99.09 100.56 99.61 99.85 99.90 100.78 100.12 99.75 99.04 99.05 99.84
Ba 222 109 170 503 1692 910 105 1116 1081 295 406 574
Ce 29 14 32 62 399 189 264 151 180 41 40 58
Cr < 2 124 6 < 2 50 14 < 2 < 2 < 2 20 226 289
Ga 25 19 23 20 25 22 30 26 27 21 24 22
La 15 < 4 16 31 217 103 163 93 102 19 21 38
Nb 2 < 2 6 14 15 12 122 56 66 2 19 5
Ni 7 94 12 4 42 31 3 4 7 20 101 189
Pb 14 < 3 < 3 3 21 6 19 4 5 < 3 < 3 7
Sc 35 31 37 23 12 20 2 13 17 11 25 26
Th 3 < 2 3 6 37 18 27 10 10 3 3 2
V 242 187 276 8 138 183 < 3 84 156 64 180 164
Y 34 14 27 49 29 30 54 33 33 7 23 17
Zr 87 33 112 248 435 314 910 351 408 99 125 99
Analysis by XRF except FeO (titration) and LO1 (loss on ignition). Lower levels of detection (ppm) for trace elements: Ba, 8; Ce, 8;
Cr, 2; Ga, 2; La, 4; Nb, 2; Ni, 2; Pb, 3; Sc, 2; Th, 2; V, 3; Y, 2; Zr, 3.
Dolerites: 1 = dolerite, Shimane Peninsula; 64 = gabbro, Shimane Peninsula. Ushikiri Formation: 43 = basalt, Shimane Peninsula;
47= rhyolite, Shimane Peninsula. Matsue Formation: 17= basalt, Matsue City; 18 = basalt, Matsue City. Oki Dozen Volcanics:
50 = trachyte dike, Oki Dozen Island; 51 = basalt, Oki Dozen Island; 56 = basalt, Oki Dozen Island. 24 = Wakurayama Andesite.
15 = Daikon Jima Basalt. 37 = Yasugi Basalt.
A 12 I)50
Na20 iI ~1 + 051 Alkaline
K20 18 e56 7 Subslk=Zline e47
t¢ e
, L , , , , i , , , , , i , , i
44 5o ss s2 s5 Si02 74
!I' FeO 43,,, ~
MgO
e24
s3 Si02 e3
Fig. 2. Major element oxide discriminant diagrams. Analyses
recalculated to 100% on an anhydrous basis. Refer to Fig. 1
and Table 1 for sample locations and ages. A. Total alkalies vs.
SiO 2 (wt.%). Discriminant line is from Miyashiro [51]. B.
FeO/MgO (i.e. all Fe as FeO) vs. SiO 2 (wt.%). Discriminant
line is from Miyashiro [52]. Rhyolite 47 omitted.
Daikon Jima, and 37: Yasugi) plot equivocally
close to the discriminant line. Of the subalkaline
samples (excluding rhyolite 47), all are tholeiitic in
terms of FeO/MgO (Fig. 2B) apart from the
Wakurayama Andesite.
5.2. Trace element chemistry
Trace dement data in this study are confined to
that available by XRF, apart from Rb, Sr, Sm,
and Nd which have been measured by isotope
dilution. Of the dements available (Tables 1 and
2), we identify three broad groupings. The first
group contains dements with high bulk distribu-
tion coefficients (i.e. D >> 1) such as Ni, Cr, and
V, termed compatible dements. The remaining
dements are those with either D << 0.1 or D < 0.1
(Rb, K, Nb, Ba, La, Ce, Zr, Sr, P, Ti, Y) which are
divided into large ion lithophile (LIL) elements
(e.g. Rb, K, LREE) and high field strength (HFS)
elements (e.g. Ti, Zr, Nb). Three MORB-normal-
ised spidergrams (Fig. 3A-C) summarise the LIL
and HFS elements. For comparison, we include
element plots from southwest Pacific, an average
island arc tholeiite (arc environments), and an
average ocean island basalt (within plate).
Figure 3A. Dolerites (e.g. 1, Fig. 3A) resemble arc
basalts, which are characterised by high Ba/Nb 33,0
and Ce/Y. 64 has lower LIL and HFS (and higher
Ni and Cr) and is interpreted as being less frac-
tionated than 1. Ushikiri Formation basalt 43 has
higher Rb/Sr and lower Ba/Nb.
Figure 3B. For the interval Sr-Nb, Matsue For-
marion basalts 17 and 18, and the Wakurayama
Andesite 24 all show arc signatures (cf. Fig. 3A,
line B). For the Ce-Y interval, the Wakurayama
Andesite plots at lower levels than line B of Fig.
3A. We note that for the same interval, samples 17
100
50
10
5
0.1
I I I I I J I I I I
100
m 50
n-
O
E m
1
0.1 / ....... /X..~ /',
• / ~ ",,~ s t ..... ~-
• ~'.. /~ ? ....... ', 17 X \~///\ \ ............... 18 ~,
\N\x~/// \\~7 \ ..... ...&..
24 I I I I I I I J I B
IO0
...< 5 L.;~:-.. /'. ~ ........ -::: ....
10 ~~~'x. x
15
O. I I I I J I I I I I I
Sr K Rb Ba Nb Ce P Zr Sm Ti Y
Fig. 3. MORB-normalised spidergrams. Break in line indicates
element below detection level. Normalising values from Pearce
et al. [53]. A. Dolerites and Ushikiri Formation. B. Matsue
Formation, Wakurayam Andesite, Yasugi Basalt. C. Oki Dozen
Volcanics, Daikon Jima Basalt. In A, line B represents calc-al-
kalic basalt ( < 52% SiO2) from southwest Pacific [54]; line C,
island arc tholeiite [55]. In C, solid line (OIB) represents ocean
island basalt [55].
340
TABLE 2
Isotope data for twelve analysed samples, with K/Ar ages [3].
No. Rb Sr S7Rb STSr/86Sr Csr(T )
(ppm) (ppm) 86 Sr ( + 2 o) Detailed analytical procedures discussed in Appendix 1
Sm Nd laTSm 1"3 Nd/l'~lNd CNd(T ) Age a
(ppm) (ppm) 144N d (+20) (Ma)
1 11.92 354.9 0.0971 0.704537+19 +0.48 4.52 16.71 0.1637 0.512788+20 +2.95 14.2
64 5.30 258.8 0.0593 0.704958 5:49 + 6.56 1.75 5.90 0.1807 0.512782 5:25 + 2.81 12.9
43 30.18 89.25 0.9780 0.704398 + 14 - 3.86 3.85 16.35 0.1432 0.512759 5:17 + 2.40 13.4
47 48.34 277.5 0.5039 0.7044945:24 - 1.19 6.95 30.80 0.1374 0.512783 5:26 + 2.89 13.1
17 101.0 1401 0.2085 0.705963 5:20 + 20.48 18.97 138.9 0.0831 0.512471 5:18 - 3.14 11.2
18 55.51 857.8 0.1817 0.704988 5:16 + 6.71 7.14 45.9 0.0947 0.512605 _+ 26 - 0.55 10.8
24 26.34 724.9 0.1051 0.704488 5:50 - 0.18 2.00 11.71 0.1040 0.512794 5:20 + 3.06 5.0
15 29.01 416.4 0.2015 0.704740 5:34 + 3.41 4.89 21.03 0.1406 0.512653 5:17 + 0.25 0.1
37 27.48 1077 0.0738 0.705495 5:16 + 14.12 3.78 22.89 0.1005 0.512580 :k 16 - 1.15 1.2
50 191.9 29.09 19.088 0.707476 5:37 + 18.86 11.91 77.71 0.0927 0.512411 + 17 - 4.39 6.1
51 78.37 1016 0.2232 0.705629 + 24 + 15.86 10.23 60.42 0.1024 0.512314 + 26 - 6.20 6.3
56 75.96 973.9 0.2256 0.705460 5:23 + 13.44 10.03 61.94 0.0979 0.512408 5:20 - 4.45 6.0
Gnl b 110.8 344.8 0.9318 0.7331205:18 11.61 71.21 0.0985 0.511230 + 18
Gn5 b 88.49 259.4 0.9893 0.731368+20 10.76 51.14 0.1272 0.511749+19
" Morris et al. [3].
b Old Dogo gneiss [44].
Standard results. BCR-I: S7Sr/S6Sr=0.705014+27 (2o mean, n=3), Rb=46.45 ppm, Sr=327.8 ppm (n=l). NBS-987:
S7Sr/S6Sr = 0.710262+25 (n = 11). BCR-I: 143Nd/1~Nd = 0.512638 + 5 (n = 18), Sm = 6.61 ppm, Nd = 28.88 ppm (n = 1). nNdfl
[57] = 0.511900 + 8 (n = 10). Blank values < 0.3 ng Rb, < 1 ng Sr, < 0.3 ng Sm, < 1 ng Nd. c s, (T) and c Nd (T) values in Table 2
were calculated using the following bulk earth parameters: S7Sr/S6Sr (present day) = 0.7045, S7Rb/S6Sr = 0.0839, )~(S7Rb) = 1.42 ×
10-11 y-l, 143Nd/n44Nd (present day) = 0.512640, n47Sm/l~Nd = 0.1967, )~(147Sm) = 6.54× 10 -12 y-1.
and 18 show disparate patterns (e.g. Ce/Y), closer
to arc rocks than OIB (Fig. 3C). The elevation of
17 in LIL elements (note also La, Sm, and Th in
Tables 1 and 2) is accompanied by higher Ni and
Cr, thus enrichment by fractionation is not viable,
but must reflect either a smaller degree of partial
melting, a LIL-enriched source, or minor (crustal)
contamination. The apparent conflict between the
positioning of 17 and 18 in the alkaline field of
Fig. 2A, and the arc trace element pattern is
discussed later. Yasugi basalt 37 has high Ba/Nb,
and a similar Sr-Ba pattern to arc magmas.
Another Sr-Nd isotope study of Yasugi basalts [9],
also noted high Sr contents (1126-3283 ppm) and
low Rb/Sr ratios (0.007-0.014).
Figure 3C. In marked contrast to Fig. 3A and 3B,
the three remaining basalts (Old Dozen, and
Daikon Jima samples) show trace element pat-
terns strongly resembling within-plate (i.e. OIB
and continental alkalic) basalts (i.e. low Ba/Nb,
Ce/Y, Ti enrichment) as shown in Fig. 3C.
Based on LIL and HFS element patterns, al-
kalies/silica, FeO/MgO and petrography, we
identify three associations in the Shimane-Old re-
gion: (1) Arc rocks (Ushikiri Formation, Dolerites,
Wakurayama Andesite, Yasugi Basalt).
(2) Rocks with a strong OIB and continental
volcanics signature (Old Dozen Volcanics, Daikon
Jima Basalt).
(3) Rocks showing transitional arc-within-plate
patterns (Matsue Formation).
5.3. Isotope chemistry
The twelve samples analysed for 143Nd/144Nd
and 87Sr/86Sr (Table 2) are plotted in terms of
eNd- CSr in Fig. 4A, B. The fan-shaped array [35]
for island arc rocks (also including other arc set-
tings; e.g. andesites from Ecuador and Chile
[36,37]) encloses some analyses of this study, al-
though the twelve analyses plot at lower ~Nd and
have a steeper slope. Available Nd and Sr isotope
data for Japanese volcanic rocks are plotted in
Fig. 4B. Analyses from the Shimane-Oki region
overlap with those from the Setouchi area [6] and
basalts + inclusions from Old Dogo Island [8] but
are more radiogenic in terms of Sr than arc-related
basalts from central and northern Japan [4,5].
In more detail, Fig. 5 shows that Ushikiri For-
mation and Dolerites plot at depleted ~Nd levels.
The separation of dolerite 64 from dolerite 1 in-
E
Nd(T) 9 A MORB
75 ~-~Aleutian Arc
0 "e'~,~ Kerguelen
-5
-7
-30 -2'0 -1'0 I
0 1'0 2'0 3'0 4'0 5'0
E Sr(T)
9 13 7
~Central and Northern 5 ~ "~ Japan
Nd(T) 3
0
-3 Oki Oogo Island
-5
-7
a 0 I i -~o-~o-,o o ;o ~'o ~o ,'o
~- Sr(T)
Fig. 4.A, end vs. Cs, for' analyses of this study (solid circles)
and fields for other areas: MORB [56], Aleutian arc [57,58],
Kerguelen [42]. Fan-shaped lines enclose array of destructive
plate margin volcanics [35]. B. end vs. ES, for samples of this
study (solid circles) and volcanics from elsewhere in Japan:
central and northern Japan [4,5], Setouchi region [6], Oki Dogo
island [8]. 341
volves increased C Sr with no change in end. We
believe that this could reflect minor post-mag-
matic alteration, possibly due to interaction with
seawater [38,39] as these rocks have been em-
placed in a marine environment. The other subal-
kaline sample (Wakurayama Andesite, sample 24)
also plots close to the field of Dolerites and
Ushikiri Formation.
Another tight grouping of data is shown by the
three analyses from Oki Dozen Island which plot
on the higher c s, and lower c Nd side of the mantle
array. Based on available data, there is significant
separation of alkaline volcanics from Oki Dogo
Island [8] which plot at higher end than volcanics
from Oki Dozen (CN~ = 0). Rhyolite 47 is isotopi-
cally indistinguishable from basaltic counterparts
in the Ushikiri Formation and we take this as an
indication that it evolved without the influence of
continental crust.
Clear isotopic separation is shown by the two
samples from the Matsue Formation. The horn-
blende and biotite-bearing sample 17 has higher
8VSr/86Sr and lower 143Nd/144Nd than sample 18.
An earlier isotope study [9] suggested that the
Yasugi basalts represented a heterogeneous man-
tle source: basalt 37 has the highest Cs~ and lowest
4
2 ¸
0
E Nd(t)
--2"
"-4
-6 -, ,4--- Dolerites, Ushikiri Formation and
~43 ~7~01 _-_-_'----6-4_-0_ -) Wakurayama Andesite
15• Daikon Jima basalt
x~18 ~ 37 Yasugi basalt
"'-~ ~ Matsue Formation
.... _pi_~
,i~s6 ~oe;
~ 51////Oki Dozen volcanics
I I i i i i l
-8 -4 0 4 8 12 16 20
E Sr(t)
Fig. 5. ~Nd VS. ~sr for samples of this study. See Table I for locations. t
24
342
~Nd recorded from the Yasugi group. The Daikon
Jima Basalt (31) plots close to mantle-derived
spinel lherzolite from Old Dogo Island [8].
6. Summary
We make the following observations based on
age of samples and their geochemical and isotopic
characteristics:
(1) There is an overall similarity in the age and
isotopic character of the Ushikiri Formation and
Dolerites. Furthermore, both suites and the
younger Wakurayama Andesite (5 Ma) have trace
element patterns consistent with island arc
volcanics modified by low-pressure fractionation
and are derived from a depleted source.
(2) Despite the similar age and close proximity
of the two samples from the Matsue Formation,
they show distinct petrographic and chemical
characteristics. Both samples are variably enriched
in the LIL elements Ba, Sr and Rb yet have a
Ce-Y pattern more resembling arc basalt, with
distinctly different end and Csf.
(3) Basalt from Yasugi (37) has an arc trace
element signature, high CSr and low end, carries a
hydrous mineral assemblage, and is un-
fractionated in terms of Ni and Cr; a nearby
basalt of the same age and mineralogy carries
upper mantle fragments [26]. We take this as
evidence of either a source region for basaltic
rocks beneath the Shimane area which is locally
enriched in some LIL elements and 875r, or minor
contamination of an unfractionated melt by a
radiogenic and LIL-enriched component such as
continental crust.
(4) In marked contrast to observations (1) to
(3), Oki Dozen and Daikon Jima basalts have
within-plate alkalic trace element patterns, but
variable isotope characteristics: 15 plots close to
bulk earth, whereas 56 and 51 extend to lower end
and higher CSr along the mantle array. Some con-
tinental basalts, and alkaline basalts from the
Kerguelen Islands have the same ~Sr-~Nd as Oki
Dozen Island [40-42].
(5) Nd and Sr isotope variations with time for
volcanic rocks occurring over a limited area of
northwest Honshu have been previously reported
[5]. This study suggested that increasing end and
decreasing CSr with time was due to progressive
insulation of the magma conduits with time lead- ing to magmas less radiogenic in terms of Sr. We
do not see such a well-defined relationship for the
rocks of the Shimane Peninsula-Matsue area and
can therefore only assume that they are tapping
different magma sources.
There is a significant spread in isotopic data
from the basaltic rocks examined here from the
Shimane region, but identifying contributing com-
ponents (e.g. subducted plate, mantle wedge, sedi-
ment, crust) by examining the chemistry of
volcanic products is hampered by the overprint of
low-pressure fractionation, and the lack of some
data (e.g. Pb and O isotopes). However, one suite
of samples (Yasugi basalts) contains upper mantle
xenoliths, elevated Ni and Cr (i.e. not
fractionated), high alkaline earth contents, and
contains modal biotite, indicating that one source
at least was LIL enriched and hydrous. We will
further examine this basalt, as it offers an insight
into possible source region compositions and may
also have some application to older rocks, espe-
cially the Matsue Formation basalts which have
tapped one source showing similar enrichment
patterns to the Yasugi Basalt, and another which
is less LIL enriched and has a different isotopic
signature.
High Ni and Cr, and LIL element abundances
in 37 cannot reflect small degrees of partial melt-
ing, as the HFS elements (e.g. Zr and Nb) are also
incompatible during partial melting (if not re-
tained in an HFS-bearing residual phase), yet are
not enriched in 37. Instead, it is likely that high
LIL and compatible element contents are due to
melting of an unfractionated source selectively
enriched in LIL elements. Such enrichment could
be due to (a) interaction of downgoing plate (pre-
sumably the PSP) with overlying mantle wedge
(by either aqueous fluid transfer or hybridisation
of slab-derived melts), or (b) crustal contamina-
tion [43-46]. Two points favour fluid transfer:
(1) Fluid release into the overlying mantle wedge
is the most viable process at depths < 100 km
[43,45]. This is consistent with the PSP being
approximately 70 km beneath the Shimane area
[10,16].
(2) Experimental work and observations of nat-
ural assemblages [45] has shown that the fluid
derived by slab dehydration would have a high
LIL/HFS ratio, inherited by any subsequent melt.
Such a pattern is preserved in the Yasugi basalts.
Rapid ascent times of magma (precluding signifi-
cant fractionation and/or contamination) are in-
dicated by included ultramafic xenoliths. Sporadic
and patchy fluid transfer could produce a hetero-
geneous source region capable of yielding closely
spaced basalts of differing isotopic and trace ele-
ment chemistry (e.g. Matsue Formation; Daikon
Jima and Yasugi Basalts).
According to seismic refraction studies under-
taken in the San'in district [47], the lower crust is
well developed and must be considered as a poten-
tial contaminant. Basic granulite from Old Dogo
Island has been studied in terms of its Nd and Sr
isotope chemistry [48], yielding values of 87Sr/86Sr
= 0.7070 and 143Nd/144Nd = 0.51245. Values for
the constituent minerals plagioclase (Sr= 1530
ppm; Nd -- 0.9 ppm), clinopyroxene (Sr = 58 ppm;
Nd = 6.3 ppm), and orthopyroxene (Sr = 23 ppm;
Nd = 1.0 ppm) are similar to values for the bulk
granulite (Sr = 717 ppm; Nd = 3.2 ppm). We stress
that the Nd values of both the constituent miner-
als and the bulk rock are significantly lower than
that of dolerite 1 (16.7 ppm). Mixing calculations
between dolerite and granulite (or its constituent
minerals) produces a concave-up curve as opposed
to a concave-down curve that would be required if
basalts 17 and 18 were mixing products. In order
to explain 17 and 18 by mixing, we must invoke
the difficult steps of either selective migration of
Nd from the basic granulite, or selective elimina-
tion of Nd from dolerite 1. Furthermore, a high
degree of fractionation would be required to ex-
plain the high Nd contents of samples 17, 18 and
37; the high Ni and Cr and the presence of
high-pressure xenoliths for 37 clearly does not
support extensive fractionation. As a result, we do
not favour crustal contamination as a viable model.
The Old Islands occupy a spatially separate
position set back from the Shimane region, where
the subducting Pacific plate is deeper (approxi-
mately 400 km according to fig. 2 of Kagami et al.
[8]) and the PSP has no influence. Some workers
have maintained that the trace element signature
of Southwest Japan aikalic rocks combines the
weak imprint of an arc contaminant on a mantle
plume [32]. Spidergrams of Oki Dozen volcanics
are similar to OIB and continental volcanics, and
their isotope characteristics resemble continental
volcanics and strongly overlap with Kerguelen
Islands volcanics, derived from an enriched man- 343
tie plume [42]. We believe that trace element
chemistry, isotopes and the back-arc setting can
best be resolved in terms of an enriched mantle
plume, analagous to the Kerguelen Islands. One
contaminant candidate is arc material [32].
7. Implications for opening of the Japan Sea
Heat flow, magnetic anomalies and sediment
distribution accord with an extensional origin for
the Japan Sea [14] although this back-arc basin is
unusual in that it has developed adjacent to a
continental margin and has (at least in part) a
floor composed of continental crust. An analagous
situation is Bransfield Strait [49] where volcanic
activity is associated with the onset of back-arc
spreading . Here, a wide range of tholeiitic, calc-
alkaline and mildly alkaline rocks has been de-
rived from a common source that has undergone
variable degrees of partial melting and fluid con-
tamination from the downgoing slab. In a review
of back-arc volcanism, Saunders and Tarney [50]
have argued that a range from N-MORB to island
arc basalts can be produced by flnid-mantle inter-
action. Taking these observations into account
and the Miocene plate tectonic situation in South-
west Japan, we maintain that the initiation of PSP
subduction, a change in the regional stress field at
15 Ma [13], tholeiitic volcanic activity in South-
west Japan, and the opening of the Japan Sea are
all inter-related.
8. Conclusions
We present the following model for the Miocene
to Recent igneous evolution of the Shimane-Oki
region:
(1) Separation of Southwest Japan from the
Asian continent occurred at about 15 Ma and was
accomplished in < 1 Ma [21]. Coincident with this
was a change in the regional stress field [13] and a
period of tholeiitic volcanism [3] the Ushikiri For-
mation and Dolerites discussed here. This
volcanism coincided with PSP subduction follow-
ing opening of the Shikoku Basin, and represented
melting of a depleted source (Fig. 6A).
(2) The Matsue Formation represents melting
of a heterogeneous source, probably the mantle
wedge overlying the subducting plate. Transfer of
fluid from the plate to the wedge provided a fertile
344
A
Ushikir i Form~
Esr ~- Nd
tion ancl Dolerites
seawater alteration
e64
\ Dolerites and
Ushikiri Formation
(Middle Miocene)
Matsue Formation
depleted source (Middle - Late Miocene)
%
component from
\subducted slab
C Oki Dozen Volcanics,
Jrayama Andesite Wakurayama Andesite,
15 Daikon Jima and
~, 037 and Oki Dogo volcanics Yasugi basalts
• ~ (Pliocene - Recent)
"', Oki Dozen
\enriched mantle component \
Fig. 6. Cartoon depicting three stages in igneous activity of
Shimane-Oki region. A. Depleted source and possible seawater
alteration for Miocene Ushikiri Formation and Dolerites. B.
Combined depleted source and enriched, slab-derived compo-
nent yields Matsue Formation. C. Variability in Oki Dozen
alkaline volcanics and Daikon Jima Basalt explained by varia-
ble enrichment of mantle plume. Separate source for Oki Dogo
and Yasugi Basalt.
source region for the production of transitional
alkalic arc magmas of differing composition. The
fluid may also have initiated melting by lowering
the peridotite solidus (Fig. 6B).
(3) Further back off the plate, alkaline
volcanism began on Oki Dozen Island at ap-
proximately 6 Ma [3] and switched to Oki Dogo at
about 3 Ma [34]. These volcanics have high al-
kaline earths and isotopes indicative of an en-
riched source and represent an enriched mantle
plume (Fig. 6C). Subduction-related magmatism
involving melting of a depleted source isotopically
similar to that of middle Miocene volcanics, is
recorded in the Miocene-Pliocene of the Shimane
area by the Wakurayama Andesite. (4) Localised heterogeneities in the Shimane
area are preserved in volcanics of < 1 Ma, where
the Yasugi Basalt carries high-pressure xenoliths,
has high Ni and Cr, LIL element contents yet an
arc trace element signature. In contrast, the
youngest lava examined (Daikon Jima Basalt) has
a pronounced alkalic signature.
Acknowledgements
P.A.M. thanks H. Honma and staff at ISEI for
the opportunity to measure isotopes, and for their
hospitality. His work was carried out under a
JSPS scholarship. Thanks also to staff of the Geol-
ogy Department, Shimane University especially T.
Watanabe who made valuable comments on the
manuscript and provided unpublished data. E.C.
Leitch read an early draft of this paper. Two
anonymous reviewers provided worthwile com-
ments.
Appendix 1--Analytical procedure for Sr and Nd
isotopes
Depending upon Sr and Nd concentrations, between 100
and 250 mg of sample was dissolved in a mixture of 40% HF (2
ml), 65% HNO 3 (0.5 ml) and 30% HC1 (0.5 ml) in a Krogh-type
teflon vessel. The vessel was loaded into a self-sealing stainless
steel jacket and heated at 210°C for 7 days. Rb, St, Sm and
Nd were separated following the procedure of Kagami et al.
[59]. REE remaining in the ion-exchange resin after Sr sep-
aration were extracted by 6 M HC1. Sm and Nd were separated
from other REE in another column using cation exchange resin
(Dowex AG 50W-X8, 200-440 mesh, H-form) by 0.2 M hy-
droxy-alpha-isobutyric acid adjusted to pH = 4.5.
The extracted dements were loaded onto a Ta-filament in a
double filament mode. Mass spectrometric analyses were made
using a MAT 261 mass spectrometer. STSR/86Sr ratios were
normalised to 86Sr/aSSr = 0.1194 and 143Nd/la4Nd ratios were
normalised to 146Nd/ln4Nd = 0.7219. S7Sr/86Sr and
143Nd/144Nd ratios were obtained by measurement of 200 and
400 scans respectively. The blank for the whole procedure was
<0.3ngRb, <lngSr, <0.3ngSmand <lngNd.
Sr isotopic ratios for NBS987 and BCR-1 were measured 11
and 3 times respectively. The averages of these ratios were
0.710262+0.000025 (20 mean) and 0.705014+0.000027 (20
mean) respectively. The average error for each sample was
0.0037% at the 20 level.
Rb and Sr concentrations were obtained by a STRb-S6Sr
mixed spike. Values of BCR-1 were 46.45 ppm Rb and 327.8
ppm Sr. We estimate an error of 1% for the Rb/Sr ratio based
on data reproducibility.
Average 143Nd/144Nd for BCR-1 (18 separate runs) and
Caltech standard nNdfl (10 runs) was 0.512638 + 0.000005 (2 o
mean) and 0.511900+0.000008 (20 mean) respectively. The
error for each sample was 0.0040% at the 2 o level. Sm and Nd
concentrations of each sample were determined using a 149Sm-
15°Nd mixed spike. The results for BCR-1 were 6.61 ppm Sm
and 28.88 ppm Nd. We estimate an error of 1% for the Sm/Nd
ratio of each sample based on data reproducibility.
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|
Morris 1989 NdSr isotope volcanic rocks SW Japan.txt
|
Magnetic dipole anomalies as indicators of mantle wedge
serpentinization
Yukari Kido
Institute for Research on Earth Evolution, Independent Administrative Institution, Japan Agency for Marine-Earth
Science and Technology (JAMSTEC), 3173-25 Showa-machi, Kanazawa-ku, Yokohama City 236-0001, Japan(kidoy@jamstec.go.jp)
Motoyuki Kido
Research Center for Inland Seas, Kobe University, Kobe, Japan
Now at Research Center for Prediction of Earthquakes and Volcanic Eruptions, Graduate School of Science, Tohoku
University, Sendai 980-8578, Japan (kido@aob.geophys.tohoku.ac.jp)
Kantaro Fujioka
Observation and Research Department, Global Ocean Development Inc., 13-8, Kamioookanishi 1 choume, Yokohama
City 233-0002, Japan (fujioka@godi.co.jp)
[1]Fine-scale magnetic anomaly data based on high-density airborne surveys have revealed several
dipole anomalies along the Chichibu Zone (CZ), which lies parallel to the Nankai Trough and forms a
part of the accretionary complex of southwest Japan. A plausible explanation of magnetic sources for
the anomalies is a series of fossil serpentine diapirs involved during the accretion process. Serpentinediapirs are known to be developed in several forearc regions as mantle wedge serpentinized material
associated with dehydration of the subducting oceanic slab. Surface geological and paleomagnetic
evidence also suggests the existence of serpentine bodies in the CZ. We applied a magnetic inversion
to each of the dipole anomalies and determined the magnetic bodies in a triaxial ellipsoid
approximation. Magnetic bodies are interpreted to lie nearly parallel to the CZ and are inclined
southward. Intensities of magnetization are compatible with those expected by measured
susceptibilities of samples in the subareal CZ, and their directions are roughly the same as thecurrent geomagnetic field, implying that the induced magnetization is dominant rather than the
remanent component.
Components: 6985 words, 6 figures, 2 tables .
Keywords: forearc; GA inversion; magnetic dipole anomaly; paleomagnetism; potential field analyses; subduction zone.
Index Terms: 1517 Geomagnetism and Paleomagnetism: Magnetic anomaly modeling; 3260 Mathematical Geophysics:
Inverse theory.
Received 23 January 2004; Revised 27 May 2004; Accepted 1 July 2004; Published 25 August 2004.
Kido, Y., M. Kido, and K. Fujioka (2004), Magnetic dipole anomalies as indicators of mantle wedge serpentinization,
Geochem. Geophys. Geosyst. ,5, Q08J13, doi:10.1029/2004GC000697.
————————————
Theme: Trench to Subarc: Diagenetic and Metamorphic Mass Flux in Subduction Zones
Guest Editors: Grey Bebout, Jonathan Martin, and Tim ElliotG3G3
Geochemistry
Geophysics
Geosystems
Published by AGU and the Geochemical SocietyAN ELECTRONIC JOURNAL OF THE EARTH SCIENCESGeochemistry
Geophysics
GeosystemsArticle
Volume 5 , Number 8
25 August 2004
Q08J13, doi:10.1029/2004GC000697
ISSN: 1525-2027
Copyright 2004 by the American Geophysical Union 1 of 12
1. Introduction
[2] Recent surveys have revealed seafloor out-
crops of serpentinized material between trenches
and their associated volcanic fronts [e.g., Fryer et
al., 1992]. These materials are products of shal-
low wedge mantle that reacted with aqueous
fluids supplied from the underlying subductingslab [e.g., Tatsumi , 1989] and are thought to have
ascended to the surface as diapirs [e.g., Fryer and
Fryer , 1987; Maekawa et al. , 1993] because of
their lower bulk density [ Toft et al. , 1990]. Ser-
pentine outcrops have also been reported along
some suture zones, interpreted to be paleo-sub-
duction zones that display evidence of ultrahigh-pressure [ G u i l l o te ta l . , 2000, 2001]. Samples
from these zones exhibit high-grade metamor-
phic recrystallization, and a depleted mantle-
wedge origin is indicated by their trace element
chemistry [ Guillot et al. , 2001]. Subducting sedi-
ments and altered oceanic crust progressively
release water. Such fluids may hydrate overlyingmantle wedge as has been interpreted for the
serpentine seamounts in the Izu-Bonin Mariana
(IBM) forearc area [e.g., Fryer et al. , 1999]. In
IBM regions, these serpentine diapirs form a
chain of small seamounts parallel to the trench[Fryer et al. , 1992]. Because the serpentinization
process accompanies the crystallization of magne-
tite [ O’Hanley , 1996], serpentine diapirs are
expected to have relatively high magnetic suscep-
tibilities [ Toft et al. , 1990; Nazarova , 1994]. This
enables us to observe them as a series of magnetic
dipole anomalies.
[
3] Southwest Japan is an ideal area for such
observations, because of the availability of high-
quality fine-scale airborne magnetic data [ Geolog-
ical Survey of Japan (GSJ)and Coordinating
CCOP , 1996] and the presence of no significant
magnetic source other than the diapirs. Actually,
we can clearly observe series of distinct magneticdipole anomalies in this region. In this paper, first
we estimate morphology of the serpentine diapirs
by applying a magnetic inversion to each of the
observed dipole anomalies. Then we measure
paleomagnetic properties of rocks, which wesampled in this region. Combining these results,
we configure the nature of the serpentine diapirs
and discuss the tectonic framework of this region.
2. Tectonic and Geological Setting
[4] The southwestern portion of the Japan arc is
principally composed of several east-west trend-ing zones of old accretionary complexes. These
accretional zones are clearly bounded by the
Nankai Trough to the south, where the Philippine
Sea plate is subducting from the south, and by the
Median Tectonic Line (MTL) to the north, where
Miocene and Quaternary volcanoes have grown
(Figure 1a). These complexes become progres-sively younger as approaching to the Nankai
Trough [e.g., Taira et al. , 1992]. The band formed
by the Chichibu Zones (CZ) is one of the most
distinct tectonic zones in the region (Figure 1a). A
number of models for its formation of the com-plexes have been proposed [e.g., Maruyama et
al., 1984; Yoshikura , 1985; Taira et al. , 1989;
Yoshikura et al. , 1990], and these are still being
debated. Tsuchiya [1982] mentioned that serpen-
tine-derived sandstones and conglomerates occur
in lower Cretaceous rocks of the Kurosegawa
Tectonic Zone (KTZ) which has been specifiedas a serpentine me ´lange in the south edge of the
CZ [e.g., Ichikawa et al. , 1956]. These rocks
exhibit sedimentary structures and consist mainly
of fragments of serpentine, together with pieces of
fragments of older basement rocks [ Ichikawa ,
1975]. This detritus was supplied mostly from a
nearby large serpentine intrusive [ Ichikawa ,
1975]. Such serpentine and ultramafic rocks are
sporadically observed along southern edge of the
CZ (Figures 1a and 1b), and they seem to have
been emplaced at various times since the Creta-
ceous [e.g., Tsuchiya , 1982].
3. Magnetic Anomalies
[5] A fine-scale data grid of residual magnetic
anomalies relative to the International Geomag-
netic Reference Field (IGRF) for this region is
available from the GSJ [1996]. The data consist
mainly of airborne and shipboard observations
merged into a uniform 1 arc-minutes grid size
with the upward continuation at 3200 m on landand coastal area and at the sea-surface in offshore
region. Along the marginal zone between the two
different continuation height, data are smoothly
connected with a weighted-mean correction. Our
research area in this paper is limited on land and
hence far from this margin. The resolution of the
magnetic data is sufficient to portray the small-scale ( /C2420 km) dipole anomalies in the region of
interest, because the airborne survey tracks exist
for almost every 30 arc-seconds.
[
6] The magnetic anomaly map (Figure 1b) clearly
shows two areas of strong signals. One is associ-ated with Miocene and Quaternary volcanoes
Geochemistry
GeophysicsGeosystemsG3G3 kido et al.: magnetic dipole anomalies 10.1029/2004GC000697
2o f1 2
15252027, 2004, 8, Downloaded from https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2004GC000697 by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
north of the MTL, and the other is related to
the magnetic lineations on the oceanic crust tothe south. It should be noted that the magnetic
lineations on the oceanic plate can still be observed
after having been subducted beyond the trench
axis (Nankai Trough). Wedged between these
two areas, magnetic anomalies are flat in theaccretionary zone except for several dipoles along
the CZ, which are indicated by boxes in Figure 1b.
Dimensions of these dipole anomalies are very
small (10–20 km) and their amplitudes are also
small compared to those of volcanic origin. How-
ever, one can observe systematic distribution of
such anomalies using the high-resolution magneticdata and with an eidetic of the well illustrated
geological map.
4. Magnetic Source Inversion
[7] To illustrate the buried magnetic sources, we
conducted magnetic inversions to the observed
magnetic anomalies. We examined the areasenclosed by bold Boxes 2, 3, 4, and 7 in Figure 1b,
where the dipole anomalies can be seen clearly.
Considering the difficulty or non-uniqueness in
the inversion of potential field, we omitted other
Boxes where pattern of the anomalies are not
simple. Enlarged views of the magnetic anomalies
Figure 1. (a) Geological map of SW Japan [ GSJ, 1995] for the area enclosed by a rectangle in the regional
map to the right. Tectonic lines are developed parallel to the trench, which is a typical feature in accretionarycomplexes. They are the Median Tectonic Line (MTL) shown by a red line, the Chichibu Zone (CZ) shown by agray zone in the middle of the Shikoku Island, the Kurosegawa Tectonic Zone (KTZ) at the southern portion ofCZ, and Nankai Trough shown by a black jagged line. Color indexes represent ultramafic, sedimentary rock,pluton, volcanic rock, and violet zone areas of serpentine research sites. (b) Magnetic anomaly map [ GSJ, 1996]
for SW Japan showing the area and map projection corresponding with Figure 1a. The data have 1
0uniform grid.
The magnetic anomalies are relative to the IGRF and are coordinated with an upward continuation at 3200 m onland and coastal area and at 0 m on the ocean. Several magnetic dipole anomalies, indicated by boxes, arerecognized along the Chichibu Zone (CZ). We applied magnetic inversion in Boxes 2, 3, 4, and 7, enclosed bythick lines, where shapes of the dipole anomalies are simple. It should be noted that no topographic signalscorresponding to the dipoles are observed.
Geochemistry
GeophysicsGeosystemsG3G3 kido et al.: magnetic dipole anomalies 10.1029/2004GC000697 kido et al.: magnetic dipole anomalies 10.1029/2004GC000697
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in the selected areas are shown in the top row in
Figure 2. Since the pattern of the dipole is very
simple, the magnetic source should be one blob(s),
except for Box 3 where two blobs are expected.We assumed a uniformly magnetized triaxial ellip-
soid to approximate the blob, which should be a
better representation of the shape of a diapiric
source with a smaller number of free parameters
than a fourth facet of the polyhedron widely used
in this type of analysis. Considering the non-
uniqueness of magnetic inversions, the number offree parameters describing a magnetic source
should be as small as possible. The ellipsoid
employed in the inversion is described by
11 parameters at most (22 parameters for Box 3).
As listed in the left column of Table 1, these are:thex-y-zposition of the center of the ellipsoid ( Xe,
Ye,Ze), the lengths of three axes of the ellipsoid
(Ae,Be,Ce), the dip angle ( D) and strike ( S)o ft h e
ellipsoid, the inclination ( Im) and declination ( Dm)
of magnetization, and the intensity of magnetiza-
tion ( m). Synthetic magnetic vector field Biscalculated as a integral of magnetic fields for
infinitesimal vector magnetic dipoles mover the
volume Vwithin the ellipsoid as follows:
B¼Z
m0=4p ðÞ j mj=jrj3/C16/C17
3 cos qr/C0m ½/C138 dV;
where m0is the magnetic permeability of free
space, ris a vector observe point relative to dipole
position, and qis an angle between randm. Note
that length of the dipole is equivalent to the
magnetic intensity (| m| = m). The infinitesimal
magnetic dipoles are discretized every 500 m gridfor computational convenience. Synthetic magnetic
anomaly dBis a scalar field defined as a difference
of absolute value of total field relative to that of the
IGRF as
dB¼jBþBIGRFj/C0jBIGRFj:
Thus calculated synthetic magnetic field at 3200 m
height is compared to the observation. Areas
masked with gray shade shown in Figure 2b are
Figure 2. (a) An enlarged view of the magnetic dipole anomalies in Boxes 2, 3, 4, and 7 in Figure 1b. Long-
wavelength components (>75 km) are removed. Color scales are adjusted to each box. (b) Correspondingsynthetic magnetic anomalies for the best fit models obtained individual inversions. RMS misfits of theinversions over the unmasked area are indicated. The corresponding ellipsoid projected to the surface is alsodrawn for visual comparison. Note that ellipsoids are buried close to the negative magnetic anomalies rather thanthe center of the dipole. (c) Projections of the best fit ellipsoid onto the three orthogonal planes, one of which isset perpendicular to the surface and strike of the ellipsoid. The cross point of the N–S and E–W arrowscoincides with the center of the corresponding magnetic anomaly map (Figures 2a and 2b). The definitions of theparameters for the ellipsoid are indicated in the projections of Box 4. Note that Dand Sare the negative
direction in this example.
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far enough from the dipoles and excluded from the
comparison for misfit.
[8] We used the Genetic Algorithm (GA) inversion
technique [e.g., Kido et al. , 1998] to minimize the
misfit between the synthetic and observed magnetic
fields. The essence of the GA is analogous to the
evolutionary theory of life, where a model having a
smaller misfit is assigned a higher probability to beendowed with the right to leave offspring into the
next generation of models through iterations of the
inversion. The iteration process is also an analogue
to the cross fertilization of genes in two selected
models among others, whose parameters are ex-panded into a chromosome map. Using this proce-
dure, the GA is extremely efficient at finding the
global minimum and a possible family of solutions
in a model space compared to the classical Monte
Carlo inversion, which walks around a model space
with an even probability. An advantage of the GA is
that it does not require an explicit initial model,which implies we do not need any a priori knowl-
edge of a plausible solution. In other words, GA is
free from the researcher’s prejudice. In addition,
technically speaking, the programming work is very
simple since it uses only synthetic forward calcu-lations. In this paper, we exactly follow the same
algorithm and tuning parameters as described by
Kido et al. [1998], which employ 100 parents
(models) for 100 generations (iterations). Then
totally 10,000 synthetic calculations were con-
ducted for each inversion, which is large enough
to make the statistical description of the ambiguitiespossible. One set of the inversion took a few hours
in a normal desktop PC.
[
9] The source parameters obtained all the inver-
sions are listed in Table 1. Root-mean squares of
misfits over the comparison area ( dM)a r ea l s o
indicated at the bottom of the table. The best fit
ellipsoid in each area generates the magnetic anom-
aly field shown in the middle row in Figure 2, andprojections of the ellipsoid onto orthogonal three
planes which are normal to axes of the ellipsoid
are in the bottom row in Figure 2. The best fit
ellipsoids for Boxes 2, 3, 4, and 7 are elongated
in west-east direction ( Ae), lying along the CZ ( S),
and is inclined southward ( D), with magnetiza-
tion ( Imand Dm) roughly parallel to that of the
IGRF ( Im
IGRF =4 5 /C176,Dm IGRF =/C07/C176). For Box 3,
two ellipsoids (up to 22 parameters) are solved at
once.
[10] We also tested a spherical representation ( Ae=
Be=Ce) of the source in Box 4. In this case Imand
Dmwere fixed to the IGRF, and mwas also fixed at
0.2 A/m as obtained by ellipsoid inversion, since m
is inversely proportional to Ae3(volume of the
sphere) for the case of a sphere. In other word,
uniformly magnetized sphere can be expressed by asingle magnetic dipole. In this case, the free param-
eters are only the x-y-zpositions ( Xe,Ye,Ze) and the
diameter of the sphere ( Ae). The result is also listed
in the next column of Box 4 in Table 1. Although
the spherical representation still scores an accept-able misfit, it can not reproduce the obliquelyTable 1. Parameter List of Solutions in the Inversions for Boxes 2, 3, 4, and 7a
Unit Box 2Box 3
WestBox 3
East Box 4Box 4
Sphere Box 7
Xe deg 132.89 133.55 133.71 134.34 134.34 136.50
Ye deg 33.48 33.60 33.65 33.88 33.88 34.47
Ze km 5.2 2.0 2.8 4.9 7.2 4.4
Ae km 10.0 11.1 9.1 9.7 8.3 9.3
Be km 5.0 2.4 2.6 5.3 - 5.0
Ce km 6.1 3.4 3.6 8.1 - 6.3
D deg /C025.6 /C012.7 /C010.3 /C033.8 - /C034.5
S deg /C01.9 /C015.4 /C020.1 /C011.3 - /C04.1
Im deg 9.9 31.1 37.2 32.7 45.0 35.0
Dm deg /C010.2 /C00.1 /C019.0 /C022.4 /C07.0 2.5
m A/m 0.17 0.18 0.19 19 0.19 0.15
dM nT 4.14 4.51 4.51 4.28 9.47 5.32
aThe bottom row represents RMS misfit dMover the data area for individual inversions. Since double ellipsoids were applied in the
inversion for Box 3, two sets of parameters are shown in Box 3. Xe,Ye, and Zeare the center positions of the ellipsoid in longitude, latitude,
and depth. Ae,Be, and Ceare the lengths of the triaxis of the ellipsoid in the E–W, N–S, and vertical directions. Dand Sare the dip
(positive in northward down-dip) and strike (clockwise from north) of the ellipsoid, while ImandDmare inclination and declination of the
magnetization; mis the magnetic intensity. For Box 4 we have examined the possibility of overparameterization of the ellipsoid by using a
simple sphere model, which has only 4 free parameters. Note that the values denoted in italics are fixed in the inversion for the sphere. Thedefinitions of Ae,Be,Ce,D, and Sare illustrated in Figure 2c.
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elongated contours of magnetic field, which is a
common feature in the observed magnetic anoma-
lies. Therefore we conclude that the ellipsoid rep-
resentation is not over parameterized, and that the
obtained parameters are valid for further interpre-
tation. Boxes 2 and 7 have very similar anomalies to
that in Box 4. In contrast, other Box 3 encompasses
two dipoles, which implies split magnetic sources.However, the basic results of the magnetic inver-
sion in each dipole, such as oblique elongation,
direction and intensity of magnetizations are rather
stable through the inversion. Therefore we expect
that the obtained such typical features of the mag-netic source are roughly applicable to the other
Boxes 1, 5, 6, and 8, except for the dimension
and multiply distribution (or complicated shape) of
the blobs.
[
11] It should be noted that there are slight trade-off
between center depth ( Ze) and vertical length ( Ce),and between volume ( Ae*Be*Ce) and magnetic
intensity ( m). This trade-off appears in the distri-
bution of possible solution of family, although the
solutions listed in Table 1 are still the best model
among the family in the individual inversion.
Furthermore, triaxial elliptic approximation is not
necessary perfect. Therefore we consider that the
obtained values of source parameters are not soprecious to quantitative interpretation and should
take qualitative essence of the result for further
interpretation.
5. Measurement of Magnetic
Susceptibility, Intensity, and
Other Properties
[12] In addition to the computational analysis, we
also conducted geophysical field investigations
along the CZ and laboratory measurements of
Figure 3. (a and b) Representative microphotograms of the serpentine sampled at the sites in Box 3. Opx and
Liz/Chrys denote Orthopyroxine and Lizardite/Chrysotile, respectively. Liz/Chrys is typically observed in low-temperature serpentine series. (c) A photograph of the handy susceptibility meter on the serpentine outcrop at thesite of Box 2. (d–f) Images of the backscattering electron microphotograph of serpentine samples obtained fromBox 4. (g) A huge serpentine outcrops in Box 2. Flow structures can be seen in the middle of the photograph.(h) A serpentine quarry at Aitani in Box 3. Boundaries between serpentine and sediment or chert zones can cleanlybe seen. They are also distinctive from the susceptibility values attached in the photograph (unit: 10
/C03SI).
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magnetic properties of rock-drill core samples
[Kido et al. , 2003]. We also took a magnetic
susceptibility meter (Figure 3c) during the inves-tigations for in situ measurements and determina-
tion of boundary rock type.
[
13] We obtained field samples of serpentine at
Boxes 2, 3, and 4 and measured their magneticproperties. Values of natural remanent magnetiza-
tion (NRM) range in 0.9 /C243.9 A/m at Box 2 and
0.3/C245.1 A/m at Box 4 as shown in Table 2. To
obtain the balance of ratio between the NRM and
induced magnetization, susceptibility was mea-sured at more than 20 sites per Boxes 2, 3, and 4
both in a laboratory and in situ. Susceptibility is a
clear indicator to distinguish serpentine and sedi-
mentary rocks. A photograph in Figure 3h is an
example of outcrop, which locates near the edge of
the serpentine tract in Box 3. The light gray rocks
are serpentine and browns are sediments. As shownby attached values in the photograph, susceptibility
changes orders of magnitude between serpentines
and adjacent sediments. Totally 80% samples of
serpentinites have high susceptibility values
(10
/C03/C2410/C02SI) [Kido et al. , 2003], which suggest
less altered. Thus the magnetic susceptibility meteris able to determine boundaries of rock types, the
volume of magnetic minerals, and the alteration
ratio of rocks. It was also a clear indicator of thedegree of serpentinization. During the field inves-
tigation, we observed massive serpentine outcrops
in all of Boxes 2, 3, and 4. A photograph in
Figure 3g is an example of outcrops in Box 2. It
should noted that serpentine flow pattern is devel-oped in the middle of the photograph. Tract of
serpentine outcrops were determined by in situ
measurement of susceptibility as mentioned above,
and found that the tracts extend to roughly several
kilometers long for all Boxes 2, 3, and 4.
[
14] A series of photographs in the left-half of
Figure 3 are examples of representative photo-
micrographs and backscattering electron micro-
photographs of serpentine thin sections takenfrom Boxes 2 and 3. In the photomicrographs
(Figures 3a and 3b), grains of Lizardite/Chryso-
tile are well observed, which is known to be
developed in serpentine at a low-temperature
condition.
[
15] Figures 3d, 3e, and 3f are the backscattering
electron microphotographs. We can observe bothTable 2. Summary of the Physical and Magnetic Properties of Rocks Sampled in Boxes 2 and 4a
SampleDensity,
g/cc Tc, /C176CMean k,
10/C03SIGeographic Coordinates
NRM ChRM
NRM, A/m MDF, mT Q Ratio Dec Inc Dec Inc
Box 2
01 2.59 594 ± 5 71.61 189 62 193 56 0.976 3.9 0.3702 2.48 589 ± 5 59.18 30 74 354 43 0.938 5.1 0.43
03 2.44 587 ± 5 36.03 /C047 13 /C034 27 1.080 2.3 0.82
04 2.35 - 44.43 /C047 19 /C039 27 1.660 2.2 1.02
05 2.62 - 138.00 118 26 134 /C059 3.910 2.1 0.77
06 2.63 - 176.90 3.910 2.0 0.60
07 2.61 - 138.10 12 41 13 39 3.780 3.2 0.7508 2.53 - 107.20 /C019 54 /C015 45 3.960 6.0 1.01
09 2.54 - 65.14 62 48 18 54 2.060 5.8 0.86
10 2.56 - 81.81 /C017 56 /C012 50 2.520 5.3 0.84
Box 4
01 2.65 569 ± 5 64.22 15 /C022 28 /C035 1.899 3.6 0.81
02 2.62 587 ± 5 150.70 275 /C022 5.127 3.7 0.93
03 2.59 588 ± 5 84.87 275 /C022 3.353 3.7 1.08
04 2.64 - 148.40 29 37 1.220 2.7 0.22
05 2.63 - 125.10 117 8 0.310 1.8 0.07
06 2.51 - 55.98 2.110 4.3 1.0307 2.64 579 ± 5 42.47 1.124 2.8 0.7208 2.64 - 117.90 /C035 32 2.030 2.0 0.47
aThe properties are density, Curie temperature, magnetic susceptibility k, declination and inclination of natural remanent magnetization
(NRM), and declination and inclination of secondary magnetization (ChRM) measured for more than 10 samples per each box. The Q ratiowas calculated using an Hvalue of 0.478 G. The Median Destructive Field (MDF) lost 80% of its natural remanent field during the stepwise
demagnetization.
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the primary magnetite crystal structure as light gray
and secondary growth as dark gray in the photo-
graph. This zoning structure has probably been
formed by the hydrothermal alteration. According
to our microscopic observation using magnetic
colloid behavior, we consider that Ti-rich titano-
magnetites exist in the inner part and magnetite inthe outer rim. This is because the accreted mag-
netic colloid particles were absorbed on outer rim
of the grain and no particle was found in the inner
part. This is also supported by the fact that the
cubic-shaped euhedral or subhedral grains areformed as can be seen in the photographs. Using
a thermomagnetization analysis, the rims of these
grains were found to be composed of monophasic
magnetite with a Curie temperature of 580 /C176C
(Figure 4), which indicates that almost pure mag-
netite is the main carrier of the magnetization.
[
16] In Table 2 we summarize physical and mag-
netic properties of the sampled serpentine rocks,
i.e., density, Curie temperature, magnetic suscepti-
bility k, direction and intensity of natural remanent
magnetization (NRM), and secondary magnetiza-
tion (ChMR), for more than 10 samples per each
Box. The Q ratio (Koenigsberger ratio), which is a
ratio of the remanent and the induced magnetisms,was calculated at two sites, Box 2 and Box 4, using
aHvalue of 0.478 G. Values of the Q ratio range
from 0.22 to1.08, whose average is 0.77. This
implies that induced magnetization is the dominant
component rather than remanent magnetization.Values of he Median Destructive Field (MDF) are
in a range of 1.8 /C246.0 mT (average: 3.5). 80% of
natural remanent field is lost during the stepwise
demagnetization. This implies that the serpentine
sampled in the vicinity of the magnetic dipole
anomalies is ‘‘soft’’ and likely to lose its original
m a g n e t i z a t i o ne v e ni nr e l a t i v e l yw e a kE a r t h ’ smagnetic field.
[
17] Figure 5 is a plot of susceptibility versus
density for all the samples. The densities of serpen-tine rocks are about 2.55 ± 0.10 g/cc as also listed in
Table 2, which is smaller than that of sedimentary
and altered rocks in vicinity and around the mag-
netic dipole anomalies. Density is a good proxy for
the degree of serpentinization, as has already been
suggested in other studies on serpentinized bodies
[e.g., Christensen , 1972; Horen et al. , 1996; Oufi et
al., 2002]. In the plot in Figure 5, no clear corre-
lation is found between susceptibility and density
for well serpentinized rocks (yellow circles). How-
ever, considering entire serpentinized rocks includ-
ing weak-serpentinization (blue diamonds), most of
the weak-serpentinized rocks have higher densities
Figure 4. (a) Thermomagnetic susceptibility profiles
of the representative samples of Box 2. (b) Sameas Figure 4a but for Box 4. Measurements were carriedout under the condition of heating and cooling rate at11/C176C/min. Stepwise procedure of thermoremanent
magnetization (TRM) measured at 150 /C176C, 250 /C176C,
350/C176C, 450 /C176C, 600 /C176C. The profile of /C0dk/dTshows a
maximum rate of TRM, which is 582 /C176C and 574 /C176Ca t
Boxes 2 and 4, respectively.
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with some exception. Accounting for widely dis-
tributed measures susceptibilities in Figure 5, we
can say that the degree of serpentinization varies
greatly within the serpentine outcrops on the mag-
netic dipole anomalies for Boxes 2, 3, and 4.
[18] Graphs in Figure 6 are the orthogonal vector
plots (Zijderveld plots) and corresponding equal
area plots of the remanent magnetization for
representative samples obtained from Boxes 2
and 4. Stepwise demagnetization was applied as
progressive alternating field starting from 2, 3, 4,
6, 8, 10, 15, 20, 30, 40, 50, 60, 80, through100 mT. The value of 80% intensity down is
shown as only 3 mT level. Both diagrams show
almost the same behavior of one component and
the unidirection.
6. Discussion
[19] Results of the magnetic inversion suggest that
the series of magnetic sources along the CZ is likely
to slant to the south ( D=/C010/C24/C035/C176) and elongate
in the west-east directions ( S=/C02/C24/C020/C176), which is
consistent with the stratal slope of CZ estimated by
seismic reflection surveys [e.g., Kurashimo et al. ,
2002; Kawamura et al. , 2003]. Obtained magneti-
zation intensities ( m= 0.15 /C240.19 A/m) are also
acceptable within the estimate based on the mea-
sured magnetic susceptibilities of the sampled
serpentine rocks in these areas when we assume
that the induced magnetization is dominant. This
assumption should reasonable, because the sam-
pled serpentine rock show characteristics of bothFigure 5. Plots of susceptibilities versus densities of obtained 55 rock samples from Boxes 2, 3, and 4. Half of the
plotted samples are serpentines (shown by yellow circles) and altered serpentines (blue diamonds); others are gabbros(red squares).
Figure 6. (a) Vector component diagrams (Zijderveld
plots) of the representative samples of Box 2.Stepwise demagnetization was carried out started from2, 3, 4, 6, 8, 10, 15, 20, 30, 40, 50, 60, 80, through100 mT. The value of 80% intensity down is shownas only 3 mT level. (b) Same as Figure 6a but forBox 4.
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strong NRM and induced magnetizations and weak
coercivity based on our measurements. During the
tectonic rotations of serpentine blocks, they should
be randomly re-oriented and immediately re-mag-
netized to the direction of the present Earth’s
magnetic field ( Ie=4 5 /C176,De=/C07/C176). Then they
should contribute to the magnetic dipole anomaliesthrough the long geological timescale. Obtained
inclinations and declinations of magnetization of
the ellipsoids ( Im=1 0 /C2437/C176,Dm=0/C24/C022/C176) also
support this interpretation. Furthermore, dimen-
sions of ellipsoids ( Ae=8/C2410 km) also agree
with the surface distribution serpentine outcrops
determined by in situ susceptibility measurements
using the handy susceptibility meter.
[
20] Given that magnetic dipole anomalies are
distributed only along the CZ, we propose two
scenarios. (1) Ascended serpentine diapirs from the
mantle wedge were trapped beneath some bound-ary and then this boundary was tectonically
uplifted to the near-surface forming the CZ
due to accretion process, such as duplex structure.
(2) Serpentine seamounts along an ancient forearc
were involved during the sedimentary accretion
that have formed the CZ, because we found several
unrevealed seamounts off central part of Shikoku[e.g., Park et al. , 1999; Kodaira et al. , 2000]. At
first, we define the mantle wedge serpentinization
process clearly. We assume that the mantle in the
overlying plate is being serpentinized. The trace
element chemistry of serpentines from the suturezone and some metamorphic belts suggest a strongly
depleted mantle wedge origin [e.g., Guillot et al. ,
2001]. For either scenario, the compatibility of the
obtained magnetization directions ( Imand Dm)
with the IGRF indicates dominance of induced
magnetization because remanent magnetization
can not keep its coherent orientation during sucha complex upwelling or accretion processes. These
samples from serpentine bodies do not individually
bear a remanent magnetization because of low
coercivity. We have considered possible relation
of subducting seamounts with magnetic dipole
anomalies at the early stage of our research on
crustal structure. However, recent compilation offocal point of small earthquakes suggests the depth
distribution of the top of the subducting slab
underneath Shikoku is 25 /C2430 km. This is too deep
to make any contribution to magnetic anomaly at
surface, in addition to the fact that temperaturemight be higher than Curie temperature at such a
depth. Even if a seamount was broken and
scratched up to the slab and then ascended to the
surface, it is also difficult to explain the fact thatthe dipole anomalies are aligned along a narrow
zone in the CZ. Small me ´lange of a fossil seamount
has only been reported in Box 4 site by Maruyama
[1981].
[
21] Looking at magnetic anomaly map (Figure 1b)
carefully, one can recognize a linear anomaly with
small along the CZ (Figure 1b). This implies thatthe CZ itself is also slightly magnetized for widely
diffused serpentine diapirs due to accretion pro-
cesses or possible transcurrent fault activity along
the CZ in the past. The occurrence of serpentiniza-
tion seems most likely to take place underneathsimilar conditions (600 /C176C, 20 kbar). Therefore
the serpentinites were formed by hydration of the
mantle wedge as a result of dewatering of the
subducting slab [ Guillot et al. , 2000]. This may
cause the closely spatial distribution along the CZ.
Moreover they might generally exist along high-
pressure metamorphic belts. There are some exam-ples of serpentinites from mantle wedges exhumed
in the Himalayas, Cuba, and Alps. They are
considered as enriched fluid solubles acting as a
lubricant [ Guillot et al. , 2000]. This may explain
the closely spatial association of serpentines and
peridotites along the CZ.
[
22] Considering other subduction zones, existence
of serpentine seamounts is apparent in the Izu-
Bonin-Mariana forearc [e.g., Fryer et al. , 1992].
However, no clear dipole anomaly has been ob-
served by shipboard magnetic surveys, because the
survey tracks are too sparse to detect small-scale
two-dimensional feature of anomalies. Even if we
had a dense magnetic anomaly data set, there is apossibility that magnetization had been lost due to
hydrothermal alteration or to accumulated serpen-
tine flow structure [ Kido et al. , 2003]. For instance
of the effect of hydrothermal alteration, Nakase et
al.[2003] found a strong positive magnetic anom-
aly above Rainbow hydrothermal site, which, un-
like many other sites on basaltic substratum, isestablished on an ultramafic outcrop. For an oppo-
site example of the effect of hydrothermal activi-
ties, basaltic samples on land in the Shimanto
region, southern portion of Shikoku, show that
the susceptibility varies up to three orders ofmagnitude, which seemed to be suffered by hydro-
thermal effects (Y. Kido et al., Regional variation
of magnetization of paleo-oceanic crust on Shi-
manto region, SW Japan, submitted to Earth and
Planetary Science Letters , 2004). Further alteration
might transform the strong magnetic minerals of
magnetite into less magnetized maghemite or trans-form to non magnetic clay mineral.
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[23] High-density airborne magnetic surveys can be
a powerful tool for delineating the distribution of
serpentine diapirs, which provide insights into the
dehydration processes of subducting oceanic plates.
Field measurements of magnetic susceptibilities are
also useful to find surface geological boundaries
efficiently and to estimate the amount of inducedmagnetization and volume of magnetic minerals.
[
24] Large-volume masses of serpentine are also
found in other regions. A brief introduction of
these examples is relevant to the present study. In
the California Coast Ranges, the New Idria body isobviously a diapir having a pipe-like shape, which
can be traced to a 10 km depth by the analysis of
observed magnetic dipole anomaly [ Jachens et al. ,
1995; Coleman , 2000]. Serpentines also occur as
tabular bodies along the faults associated with the
Coast Range Ophiolite. These serpentine bodies
seem to form weak-planes (or spots) along thefaults. Many serpentinite bodies have been thought
to originate in the alteration of ophiolite prior to
accretion when they became detached from the
lower plate. On the other hand the serpentines
lying along younger faults are thought to originate
in older ophiolite due to reactivation of older faults
being injected by serpentines.
[
25] The Santa Clara Formation was invaded by
serpentine diapirs in very recent time [ Page et al. ,
1999]. Page et al. interpret that the serpentine
bodies within the Franciscan Complex were de-
rived from peridotite wedges in the Franciscansubduction complex (now exposed as a me ´lange)
and were generally unrelated to the Coast Range
Ophiolite.
[
26] The ocean-continent boundary off the Iberian
margin (west of the Galicia bank) provides another
example of the serpentine diapir emplacement
[Boillot et al. , 1980]. No distinct magnetic anomaly
is observed in this offshore area although many
diapirs exit and even locates close to the passive
margin where no magnetic lineation is expected
due to the Cretaceous quiet zone. A scenario for
this region may be the same as that in Shikoku, i.e.,
the serpentinization occurred due to hydration
processes associated with the accretion of oceaniccrust, and then the resulting serpentinites ascended
diapirically along fractures [ Boillot et al. , 1980].
The reason why so few magnetic anomalies are
observed along the offshore Iberian and Izu-Bonin-
Mariana forearc examples may be: (1) Serpentinesin these areas were more pervasively altered to clay
minerals because their host diapirs were in contact
with seawater for longer geological periods thanthe example of Shikoku presented here. (2) Ser-
pentines in these areas were much progressed to
the degree of serpentinization, which results in
loosing the susceptibilities. This can be also found
in part of Shikoku showing low susceptibilities
with low magnetic anomaly signal. (3) Small-scale
magnetic signals are diminished at sea surfacethrough the deep depth of ocean.
[
27] Mantle wedge serpentinization and serpentine
diapir processes might therefore be global phenom-
ena. However, California Coast Ranges and the
Shikoku area, Japan, may be suitable area to studyremnants of serpentine diapirs than many margins
beneath ocean.
Acknowledgments
[28]The authors are greatly indebted to H. Ishizuka, H. Sato,
and S. Machida for their kind field guidance and manysuggestions for this work, S. Machida for her assistance withsampling, and H. Murakami and S. Yoshikura for theirvaluable comments. The authors are greatly indebted toH. Shibuya for his many suggestions, to Y. Tatsumi, N. Seama,T. Fujiwara, and T. Furuta for valuable comments, to K. H.Hattori and S. Arai for constructive discussion and muchpetrological information, and to Z. Zhong for paleomagneticmeasurements in detail and constructive comments. We thankG. Bebout, W. M. White, J. Dyment, and C. Finn for theirhelpful comments in improving the manuscript. GMT software[Wessel and Smith , 1995] was the supportive tool for the
mappings shown in this paper. IFREE provided the fundingfor the serpentine study in Shikoku, SW Japan.
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Geochem Geophys Geosyst - 2004 - Kido - Magnetic dipole anomalies as indicators of mantle wedge serpentinization.txt
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Stratigraphic correlation of Upper Cretaceous to Paleocene forearc basin
sediments in Northeast Japan: cyclic sedimentation and basin evolution
Hisao Ando *
Department of Environmental Sciences, Faculty of Science, Ibaraki University, Bunkyo 2-1-1, Mito 310-8512, Japan
Received 15 August 2000; revised 15 May 2002; accepted 30 July 2002
Abstract
Cretaceous to Paleocene sediments are distributed across a 200 km-wide and 1400 km-long belt in Northeast Japan to Sakhalin. Twenty-
six stratigraphic sections including land-based and offshore drill holes and surface sections are correlated to delineate the sedimentary history
and basin evolution of the Yezo forearc basin. Two second-order shallowing-upward cycles are recognized in Hokkaido but are obscure in
north Honshu. Shallow-marine to paralic sediments of the first cycle (early Albian to Turonian) are represented by the Mikasa Formation(late Albian to Turonian) as the western marginal facies, while the eustatic curves (e.g. Haq curves) reach the highstand maximum at the late
Cenomanian. The Campanian to Maastrichtian shallow- and non-marine facies of the second cycle (Coniacian to Maastrichtian) are
characterized by the Hakobuchi Group in Hokkaido and its correlatives in north Honshu. This large-scale shallowing may have reflecteddirectional changes of the Izanagi-Kula plate motion near the early Campanian. The uppermost Maastrichtian to Upper Paleocene is possiblyabsent everywhere in the basin. The shallow-marine to non-marine facies often show several third-order depositional sequences with upward-
coarsening facies successions. They may have been controlled by global eustasy even in an active-margin setting, since the stacking patterns
of sequences and the timing of sequence boundaries appear considerably concordant with the oscillation patterns of the Haq curves.q2002 Elsevier Science Ltd. All rights reserved.
Keywords: Yezo forearc basin; Cretaceous to Paleocene; Northeast Japan; Sequence stratigraphy; Shallow-marine sedimentation; Cyclic sedimentation; K/T
unconformity; Haq curves
1. Introduction
Compared with offshore marine sediments, shallow-
marine to non-marine sediments record facies changes andsuccessions more distinctively, reflecting local/global sea-level changes, tectonic movement and sediment supply.They are suitable for investigating cyclic sedimentation,sequence stratigraphy and sedimentary history throughgeologic time. Moreover, molluscan fauna often occurredfrom shallow-marine strata and some floral componentsfrom non-marine sediments provide paleoecological, paleo-botanical and paleobiogeographical information.
From the Pacific coast of north Honshu to Sakhalin
through central Hokkaido, Upper Cretaceous to LowerPaleocene shallow-marine to non-marine sediments as wellas offshore-marine sediments are distributed over 1400 kmlong in a north–south direction ( Fig. 1 ). They are thought
to have been deposited in an ancient forearc basin alongthe eastern margin of the paleo-Asian continent, called the
Yezo forearc basin (e.g. Kimura, 1994, 1997 ).
The major objective of this paper is to establish the high-
resolution stratigraphy of the Yezo forearc basin sediments,referring to chrono- and lithostratigraphic data of three land-based and eight offshore drill holes for oil exploration, inaddition to the land outcrop sections by previous researchersand my field survey. The next is to delineate the sedimentaryhistory and basin evolution of the Yezo forearc as one of theaspects of environmental change in Far East Asia.
2. Geologic setting
Fig. 1 is a pre-Neogene geological map and tectonic
division of Northeast Japan, the northern half of theJapanese Islands, and the southernmost part of Sakhalin. Itgenerally shows the complicated geologic structure that istypical of an active convergent margin. Though there aredifferent views on the post-Jurassic tectonic history ofNortheast Japan, it is generally agreed that the geologic
1367-9120/03/$ - see front matter q2002 Elsevier Science Ltd. All rights reserved.
PII: S1 36 7 -9 12 0 (0 2) 00 1 11 -6
Journal of Asian Earth Sciences 21 (2003) 921–935
www.elsevier.com/locate/jseaes
*Tel.:þ81-29-228-8391; fax: þ81-29-228-8405.
E-mail address: ando@mx.ibaraki.ac.jp (H. Ando).
H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 922
framework was formed by collision and accretionary
processes of two arc-trench systems, namely the paleo-Japan and paleo-Kuril Arc-Trench systems ( Niida and Kito,
1986; Minoura and Hasegawa, 1992; Otsuki, 1992; Isozaki,
1996; Kimura, 1994, 1997) . The Tanakura Tectonic Line
(TTL) is a left-lateral strike-slip fault and a large-scaletectonic boundary separating Northeast Japan from South-
west Japan. The pre-Neogene geology in Northeast Japan is
characterized by a zonal arrangement of several geotectonicprovinces on the surface, some of which were intenselymodified by post-Paleocene tectonism partly related to the
opening of the Japan Sea caused by backarc spreading
during the middle Miocene (e.g. Jolivet et al., 1994 ).
Furthermore, the Cretaceous and Paleogene are covered byNeogene and Quaternary volcanics and sediments in some
places.
North Honshu consists of three NNW-SSE trending
units, namely, the Abukuma, South and North KitakamiBelts from south to north, separated from each other by the
Hatagawa Tectonic Line (HTL) and the Hayachine Tectonic
Zone (HTZ), respectively ( Fig. 1 ). Hokkaido is subdivided
into eight belts: the Oshima, Rebun-Kabato, Sorachi-Yezo,Idon-nappu, Hidaka, Yubetsu, Tokoro and Nemuro Belts,
from west to east. Among them, the last three and the
eastern half of the fifth (Hidaka Belt) in Hokkaido areelements of the paleo-Kuril Arc ( Nanayama et al., 1993;
Kimura, 1997 ). The others and the three belts in north
Honshu form the N-S trending paleo-Japan Arc system
composed of (1) early Cretaceous magamatic arc, (2) lateCretaceous to Paleocene Yezo forearc basin and (3) lateCretaceous to Paleogene accretionary complexes.
The early Cretaceous magmatic arc is represented by
plutonic rocks widely distributed in north Honshu to westHokkaido (Oshima Belt) ( Tsuchiya and Kanisawa, 1994 ).
Volcanic rocks are also distributed along the Pacific coast of
the South and North Kitakami Belts. In the North Kitakami
and Oshima Belts, granitic rocks (130–115 Ma in radio-metric age: Imaoka et al., 1999 ) intruded into Jurassic to
early Cretaceous accretionary complexes. The adakitic
granite zone is inferred through the prominent positive
magnetic anomaly belt from off central Honshu to centralHokkaido (Kitakami and Joban magnetic belts: Finn, 1994 ;
Ishikari-Kitakami positive magnetic belt: Tsuchiya and
Kanisawa, 1994 ). This zone is generally situated along the
western margin of the basin. These rocks seem to have beenthe major sources of terrigenous siliciclastics in the basin(Nanayama et al., 1993 ).
The Cretaceous to Paleocene forearc basin-fill and
accretionary complexes are exposed across a 200 km-wideand 1400 km-long belt in central Hokkaido to Sakhalin
(Kimura, 1994, 1997) . The forearc basin-fill extends
southward widely into shelf and continental slope, and
narrowly along the Pacific coast in north Honshu. For a few
decades, this ancient forearc has been explored forhydrocarbon resources through seismic surveys and land-based and offshore drilling by oil corporations. Among
many drill holes, some penetrated into the Cretaceous to
Paleocene sediments beneath the Eocene and Miocene.Among DSDP sites only a hole (site 439 in Fig. 1 ) on the
continental slope landward of the Japan trench penetrated
the Cretaceous mudstone which was lithostratigraphically
correlated with the Yezo Supergroup ( von Huene et al.,
1982; von Huene, 1994 ). Seismic, magnetic and gravity
surveys have been carried out through DSDP ( von Huene
et al., 1982 ) and marine geological mapping by the
Geological Survey of Japan (e.g. Tamaki, 1978; Nakamura,
1990 ). These surveys and maps reveal that the Cretaceous to
Paleocene sediments unconformably underlie the Eocene
and Miocene as acoustic basements. A comprehensive study
of their stratigraphy and sedimentary history within theoverall Yezo forearc basin has not been attempted prior tothis study.
Late Cretaceous to Paleogene accretionary complexes
are represented by the Idon-nappu and Hidaka Belts incentral Hokkaido. These two belts may also extend south-ward beneath the modern forearc offshore of north Honshu,
especially beneath the lower continental slope to upper
trench slope ( Kimura, 1997 ). A large volume of the
accretionary prisms may have been tectonically erodedthrough subduction processes along the Japan Trench ( von
Huene et al., 1982; von Huene, 1994; Kimura, 1997 ).
Recently Ueda et al. (2000) documented the spatial relation
of the late Jurassic to Paleogene accretionary complexes andoverlying forearc basin-fill in the Idon-nappu Belt, Hok-
kaido.
The Upper Cretaceous to Paleocene shallow-marine to
non-marine deposits are only exposed narrowly andfragmentarily along the Pacific coast and widely in the
western part of the Sorachi-Yezo Belt ( Fig. 1 ). Offshore
marine sediments are also much preserved in the Yezoforearc basin-fill. Because the strata of the Yezo forearcbasin had not been subjected to intense diagenesis and
tectonic deformation in comparison with contemporaneous
accretionary complexes, they preserve good geohistoricalinformation on depositional facies and their successions,micro- and mega-fossils and others.
The distribution and geologic structure of the forearc
basin strata are considerably different between north
Fig. 1. Pre-Neogene geological map and tectonic subdivision of Northeast Japan. Surface geology is compiled from Kiminami et al. (1986), Niida and Kito
(1986), Wakita et al. (1992), Minoura and Hasegawa (1992), Kimura (1994, 1997) and Isozaki (1996) . The distribution of adakitic granite is referred to Finn
(1994), Tsuchiya and Kanisawa (1994) and Tsuchiya et al. (1999) . Offshore distributions of Yezo forearc basin-fill and accretionary complexes are based on
Wakita et al. (1992), Kato et al. (1996), Kimura (1997) and Ito et al. (1998) . White squares and arrows indicate the locations of surface sections and drill holes,
respectively. Numbers refer to columnar sections shown in Fig. 3 . TTL, Tanakura Tectonic Line; HTL, Hatakawa Tectonic Line; HKF, Hizume Kesennuma
Fault; HTZ, Hayachine Tectonic Zone.RH. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 923
Honshu and Hokkaido, partly because of post-Paleocene
tectonic deformation characterized by intra-arc left-lateralstrike-slip fault movements in north Honshu and arc–arccollision in Hokkaido ( Otsuki, 1992; Kimura, 1997 ). In
terms of the basin geometry, surface tectonic features andgeneral stratigraphy, the Yezo basin can be divided intotheee subbasins, the Hokkaido, Joban and Kitakamisubbasins in this paper.
3. Stratigraphic correlation of the Yezo forearc basin-fill
As already reported by Ando (1997) , major surface
stratigraphic sections of the shallow-marine and non-marinestrata in the Cretaceous Yezo forearc basin were correlatedthrough detailed stacking patterns of facies and third- tofourth-order depositional sequences. Fig. 2 is the revised
sequence stratigraphic correlation among the Futaba, Kuji
and Hakobuchi Groups and the Mikasa Formation, in
addition to the recent data on the Hakobuchi Group in theHobetsu area by my field survey. The stacking patterns ofdepositional sequences and the timing of sequence bound-aries and key surfaces bounding systems tracts appearconsiderably concordant in frequency and phase with theoscillation patterns of the Haq curves ( Haq et al., 1988 )
despite some discordance such as the Maastrichtian in theupper part of the Hakobuchi Group. Each group orformation constitutes two to seven third-order depositionalsequences less than 200 m thick. Some of the sequences alsoconsist of fourth-order sequences or parasequences usuallyseveral tens of meters thick. Lowstand deposits may becorrelated with major sequence boundaries of the Yezobasin sediments.
Because the surface sections are limited only to the
Pacific coastal areas, especially in north Honshu, integratedstratigraphic correlation over a wide area of the Yezo basinwill be possible by referring to chrono- and lithostrati-graphic data from land-based and offshore drill holes as wellas surface sections. Drilling data are based mainly oncuttings and a small amount of cores and geophysical welllogs. Though sections obtained from drilling data arestratigraphically continuous, accuracy of facies andsequence identification is usually lower than from surfacesections. Therefore, we need to achieve careful comparisonbetween surface and drilling data. Particularly, the authorexamined the biostratigraphic control, lithostratigaphy, andinterpreted sedimentary environments through cuttings,geophysical well logs and cores of each drill hole in detail(JNOC, 1974, 1985, 1986, 1991, 1992, 1995, 2000 , etc.).
Fig. 3 is the litho- and biostratigraphic correlation of a
total of 26 sedimentary successions including 15 surfacesections, three land-based and eight offshore drill holes,from the southernmost part of Sakhalin to Nakaminato incentral Honshu. It shows the stratigraphic distribution ofdepositional facies, upward-coarsening/-fining facies suc-cessions, index fossils and sequence boundaries. Agecontrol is based mainly on megafossils (ammonites and
inoceramids) for the surface, and microfossils (foraminifers,calcareous nannoplankton, pollen and dinoflagellates) forthe drill holes. Facies distributions are not so simple, andtime control by index fossils and radiometric ages is
somewhat scarce in some sections, but the figure reveals
some general aspects of the sedimentary history of the Yezoforearc basin-fill during the Albian to Paleocene.
The most remarkable difference between Hokkaido and
north Honshu ( Fig. 3 ) is the length of stratigraphic ranges
for columnar sections. This partly depends on the tectonicsetting and exposure condition of the sections. Forearc stratain Hokkaido are characterized by widely-distributed, similarstratigraphic units called the Yezo Supergroup ( Okada,
1983 ) along the meridian mountainous zone (Yezo-Sorachi
Belt; Fig. 1 ). They were steeply or vertically inclined by a
large-scale tight anticline with a westerly-inclined axis,which was formed by the collision of the paleo-Kuril Arc tothe paleo-Japan Arc. On the other hand, North Honshu ischaracterized by sporadically distributed, narrow surfaceexposures with short stratigraphic ranges along the coast,reflecting the simple seaward-dipping homocline or openfolding. Some drill holes, however, show continuous
geological columns.
3.1. Hokkaido to south Sakhalin: two second-order
upward-shallowing cycles
In Hokkaido, the Cretaceous forearc sediments called the
Yezo Supergroup have been conventionally divided into theLower, Middle and Upper Yezo Groups and the HakobuchiGroup since Matsumoto’s (1951) stratigraphic revision
following the quadripartite division by Yabe (1926) . Since
Matsumoto (1942, 1943) , many stratigraphic studies have
been carried out from several viewpoints (as reviewed inHirano et al. (1992) ). As the above four groups and the
underlying Sorachi Group of ocean floor sequence arerecently known to be mostly conformable, except somelocal discordance ( Kito et al., 1986; Kito, 1987; Takashima
and Nishi, 1999; Ueda et al., 2000 ), they may be categorized
down to formations in the sense of nomenclature. Theauthor follows Okada’s (1983) usage to avoid nominal
confusion here. There are two second-order upward-
shallowing cycles in the Yezo Supergroup, represented by
the Lower to Middle Yezo Groups (late Hauterivian toTuronian) and the Upper Yezo and Hakobuchi Groups(Coniacian to Maastrichtian) ( Okada and Matsumoto, 1971;
Ando, 1997 ). According to the stratigraphic duration of two
cycles, they seem to represent megasequences ( Haq et al.,
1988 ;Figs. 2 and 3 ).
The Lower Yezo Group (late Hauterivian to early
Albian: Takashima and Nishi, 1999 ) is mainly composed
of offshore (continental slope) mudstone and turbidites with
few megafossils. The lower half of the Middle Yezo ischaracterized by offshore (outer shelf to upper continentalslope) dark gray mudstone commonly bearing ammonitesH. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 924
Fig. 2. Sequence stratigraphic correlation among the major surface sections of the Upper Cretaceous shallow-marine to non-marine strata in the Yezo forearc basin. Refer to Figs. 1 and 3 for location of column:
Futaba (24), Kuji (17) and Hakobuchi Groups (8, 12) and the Mikasa Formation (9). PS, parasequence. Geologic time scale is based on Gradstein et al. (1995), Berggren et al. (1995) and de Graciansky et al.
(1998) .
H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 925
Fig. 3. Stratigraphic correlation of late Cretaceous to Paleocene sediments throughout the Yezo forearc basin from south Sakhalin to Northeast Japa n. Refer to Fig. 1 for location of each column. Geologic time
scale same to in Fig. 2 .Ando et al., 1995; Hata and Tsushima, 1969; Kase et al., 1984; Nagahama and Terui, 1992; Saito, 1961; Susaki and Lwasaki, 1992; Tanai et al., 1978; Toshi mistsu, 1988.H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 926
Fig. 3 ( continued )H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 927
and inoceramids within calcareous concretions. In contrast,
the upper half is dominated by alternating beds of sandstoneand mudstone including turbidite facies. Particularlyshallow-marine to paralic sediments are called the MikasaFormation, and are distributed only in the western centralpart of the Sorachi-Yezo Belt ( Fig. 3 , columns 7–10).
Though the Mikasa Formation shows some local and lateralfacies changes representing the western marginal facies ofthe Yezo basin, it is interpreted to be composed of threethird-order depositional sequences and three fourth-ordersequences or parasequences for each third-order ( Ando,
1990a,b, 1997 ;Fig. 2 ;Fig. 3 , columns 8–10). The age of the
formation ranges from late Albian to late Turonian as basedon ammonite and inoceramid biostratigraphy ( Ando,
1990a,b ).
The lower part of the second cycle is called the Upper
Yezo Group (Coniacian to early Campanian) characterizedby monotonous, massive and somewhat bioturbated mud-stone of offshore (outer shelf to upper continental slope)environments with ammonite and inoceramid fauna. Lithicsandstone rich in andesite and rhyolite fragments possiblyderived from westerly magmatic arc, conglomerates ofdebris-flow origin, and turbidite facies are locally developedand interfingered into the mudstone facies mentioned above(Ashibetsu area: Fig. 3 , column 8).
The uppermost unit of the Yezo Supergroup is the
Hakobuchi Group conformably overlying the Upper Yezo
Group in most places and disconformably in the Ashibetsu
area ( Fig. 3 , column 8). It disconformably underlies the late
Eocene Ishikari Group bearing coal measures. Comparedwith the underlying strata, shallow-marine to fluvialsandstone predominate in the Hakobuchi Group except inthe Nakatonbetsu area ( Fig. 3 , column 3). Floodplain
mudstone facies are subordinately developed in fluvialfacies. Acid tuff and tuffaceous sandstone indicating acidvolcanism in the westerly magmatic arc are common in theupper part (Maastrichtian). The Hakobuchi Group shows thecomplicated stacking patterns of third-order sequences andfourth-order sequences (or parasequences) ( Fig. 2 ;Fig. 3 ,
columns 1–6, 8, 11 and 12). It constitutes several repetitiveupward-coarsening units of offshore sandy siltstone toshoreface sandstone, associated with fluvial sandstone,mudstone and sometimes coaly beds as lowstand sediments.They can be observed in some large outcrop sections asdescribed later ( Fig. 4 : column 12 on Fig. 3 ;Fig. 5 : 11).
In many areas, the uppermost part of the Hakobuchi
Group usually reaching the Maastrichtian is overlain by adisconformity or a gentle angular unconformity andcovered by deposits younger than the Paleocene such asthe middle to upper Eocene Ishikari Group or the upperEocene to lower Oligocene Poronai Group. Lithologicalchanges are conspicuous at the unconformity.
Since Yasuda (1986) found Paleocene planktonic
foraminifers from the offshore siltstone facies of theuppermost part of the group in the Nakatonbetsu area, ithas been anticipated that detailed documentation of the K/Tboundary and its surroundings should be carried out through
magnetostratigraphy and biostratigraphy by taxa other thanforaminifers. Recently, Late Paleocene dinoflagellates andnannofossils were discovered from the upper part of theHokobuchi Group in the Oyubari area ( Suzuki et al., 1997 ;
Fig. 3 , column 11) and the Nakatonbetsu area ( Okada et al.,
1998 ;Fig. 3 , column 3), respectively. Until now, the
uppermost Maastrichtian and lower Paleocene strata have
not been biostratigraphically documented anywhere in the
Yezo basin. Recently the author found an importantunconformity in the Nakatonbetsu area, which seems tohave eroded the K/T boundary, uppermost Maastrichtianand Lower Paleocene. This interpretation is located onbiostratigraphy and an erosional surface ( Ando et al., 2001 ;
Fig. 6 ). This means that a different sedimentary cycle of the
Paleocene beneath the overlying post-early Eocene beds,exists in the uppermost part of the Hakobuchi Group in theNakatonbetsu and Oyubari areas, and the Yezo forearc basinhad continued until late Paleocene.
Kurita and Obuse (1994, 1997) reported the late
Paleocene to early Eocene dinoflagellates and pollen florafrom the Haboro Formation in the surface section of the
Haboro area ( Fig. 3 , column 6) and in a drill hole (column
2). The Haboro Formation is composed of sandstone andmudstone associated with several coaly beds. It seems tohave been deposited under nearshore and paralic environ-ments at the last stage of the Yezo basin, judging from thelithological similarity with the uppermost part of theHakobuchi Group.
3.2. North Honshu: sporadic stratal distribution
As described in Ando (1997) , shallow-marine and non-
marine facies predominate in every stratigraphic horizon ofsurface sections of the Futaba Group, the Kuji Group and itscorrelatives, in comparison with Hokkaido. Their strati-graphic ranges are limited to the Coniacian to Santonian andthe Santonian to possibly Maastrichtian, respectively(Figs. 2 and 3 ). However, lithostratigraphic data of offshore
drill holes from the geophysical well logs, cuttings andcores, show the presence of thick offshore marine mudstonefacies associated with shallow-marine and non-marinefacies ( Fig. 3 ). This facies differentiation between the
surface sections and the drill holes seems to reflect theirgeographic position within the basin. The surface sectionsconcerned here were situated in the western margin of the
basin at that time ( Fig. 1 ).
Though it is difficult to recognize the two upward-
shallowing cycles clearly from our stratigraphic data withinnorth Honshu, the predominance of fluvial-plain facies inthe Campanian and Maastrichtian is generally consistentwith the Hakobuchi Group in Hokkaido. For the Cenoma-nian to Turonian, columns 13 (off Hachinohe) and 15 (offKuji) on Fig. 3 indicate fluvial and shallow-marine facies
correlative to the Mikasa Formation in westerncentral Hokkaido. Column 20 (Kesennuma) also shows an
H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 928
upward-coarsening succession of shallow-marine facies in
the Cenomanian. In this column, the second shallow-marinefacies appear at the Upper Turonian to Coniacian above theUpper Cenomanian to Turonian offshore facies. This faciessuccession is similar to the east Ikushunbetsu sectiondescribed in Ando (1990a ,Fig. 3 , column G). In the same
way as Hokkaido, the uppermost Maastrichtian and lowerPaleocene strata as well as the K/T boundary have not beenfound and seem to have been eroded away, though timecontrol is somewhat scarce due to the lack of marine index
Fig. 4. Two upward-coarsening facies successions in the lower part of the Hakobuchi Group along the Hobetsu section (section 12 on Fig. 3 ). Vertically-
inclined strata about 80 m thick form the upper part of the first depositional sequence (DS1) and the main part of DS2.H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 929
Fig. 5. Depositional sequences of the Hakobuchi Group exposed along the Shuparo River around the Oyubari Dam in central Hokkaido (column 11 on Fig. 3 ).
Except DS1, eight DSs are well defined in overturned strata about 450 m thick. Most of DSs show upward-coarsening facies successions that are mainlyassigned as high stand systems tracts (HST).
Fig. 6. Photo and sketch of the erosional surface (sequence boundary) between the Heitaro and Oku-utsunai Formations in the Nakatonbetsu area, northHokkaido (column 3 on Fig. 3 ).H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 930
fossils (columns 13–15, 22). Geophysical logging data of
column 14 shows the presence of thick fluvial plain facies ofthe uppermost Paleocene to lower Middle Eocene correlatedto the Haboro Formation in north Hokkaido.
4. Depositional sequences showing upward-coarsening
facies successions in the Hakobuchi Group, Hokkaido
Taking the temporal and spatial distribution of facies
successions for nine surface and drill hole sections in
Hokkaido into account, the Hakobuchi Group is charac-terized by several (less than 10) third-order sequencesusually showing an upward-coarsening facies succession,though often associated with a thin upward-fining unit inthe basal part. The number of the sequences is different insections depending on the predominant facies. Especiallyin the Hobetsu area, nine fourth-order sequences (para-sequences) are observed in the upper part (column 12 onFig. 3 ). Two very well-exposed surface sections are
described below as examples of upward-coarsening facies
successions formed by progradation of inner shelf to delta
plain systems.
4.1. Hobetsu area
In the Hobetsu area, the Hakobuchi Group forms seven
third-order depositional sequences (DSs) composed oflowstand fluvial facies, transgressive upward-fining andhighstand upward-coarsening successions from inner shelfsiltstone to shoreface sandstone ( Fig. 3 , column 12). When a
low stand systems tract (LST) is lacking, the sequence isrepresented by thin upward-fining trangressive systems tract(TST) and thick upward-coarsening highstand systems tract(HST). The thicknesses of TST and HST are changeable
depending on their spatial position within a third-order
sequence.
Fig. 4 shows the lowest two DSs observable along nearly
vertically-inclined strata. A ravinement surface (RS) withtransgressive conglomerate above the upper part of DS1represented by trough cross-stratified sandstone, severalmeters of upward-fining sandstone, a maximum floodingsurface (MFS), sandy siltstone, interbedded hummockycross-stratified (HCS) sandstone and siltstone, amalgamatedHCS sandstone, and again trough cross-stratified sandstoneoccur successively. The lower part of DS2 contains some
bivalves (inoceramids and Apiotrigonia ) and ammonites.
Oysters shells and wood fragments are common in the righthand side of the outcrop ( Fig. 4 ).
4.2. Oyubari area
The second example is the Oyubari dam section, the type
locality of the Hakobuchi Group ( Fig. 3 , column 11; Fig. 5 ).
Nine complete DSs are recognized in the overturned strataexposed along the upper and lower stream sides of the dam.Some sequence boundaries are coincident with marine
flooding surfaces or RSs in the case where fluvial unitsfilling an incised valley are lacking. Recently, Suzuki et al.
(1997) found late Paleocene dinoflagellates and pollen flora
from DS9. Because of no critical lithofacies change withinthe group, the disconformity is presumed to be presentbetween DS7 and 8, in addition to the conspicuousunconformity with the Eocene coal measure Ishikari Group.
5. K/T gap: unconformity between Upper Maastrichtian
and Upper Paleocene
In Japan, the only well-documented K/T boundary
section is the Kawaruppu section in the Nemuro Group ofthe Nemuro Belt ( Saito et al., 1986 ). This group is regarded
as the paleo-Kuril forearc basin-fill. However, the K/Tboundary section has never been found in the Yezo forearc.
Recently the author found an important unconformity
correlated to this gap at the upper part of the HakobuchiGroup in the Nakatonbetsu area, northern Hokkaido ( Fig. 3 ,
column 3; Fig. 6 ).
A sharp erosional surface is observable between the
offshore mudstone of the Heitarozawa Formation and thelower shoreface to inner shelf fine sandstone of the Oku-utsunai Formation. The undulated base contains an inter-mittent lenticular layer of pebble conglomerate less than ameter thick. The overlying medium to coarse sandstoneabove a flat and slightly erosional surface yields Glycymeris
shell beds and pebbles and appears to have been depositedunder a shoreface storm-dominated environment. Followingthese lag deposits, amalgamated HCS sandstones bearingthinGlycymeris shell beds in the lower part, and bioturbated
silty sandstone and bioturbated sandy siltstone form an
upward-fining (transgressive) facies succession. This basalerosional surface of the Oku-utsunai Formation is thought tobe a sequence boundary associated with transgressive lags.
The underlying Heitarozawa Formation, though not from
this locality, contains ammonite fauna which seem torepresent the Lower Maastrichtian to the lower part of theUpper Maastrichtian in comparison with the biostrati-graphic and paleomagnetic study in Sakhalin ( Shigeta
et al., 1999; Kodama et al., 2000 ). Above the extinction
horizon of inoceramids, Pachydiscus flexuosus ,Zelandites
varuna andAnagaudryceras matsumotoi continue to occur
with Tenuipteria awajiensis which is somewhat similar to
inoceramids, but is systematically referred to as a differentstock. The author emphasizes that these are the newestammonites, not well known in Japan until now. Recently
Okada et al. (1998) found Maastrichtian nannofossils from a
different section.
The Utsunaigawa Formation overlying the Oku-utsunai
Formation is composed of offshore sandy siltstone andyields the late Paleocene nannofossils in the lower partabove Yasuda (1986)’s Paleocene foraminifer horizon
(Okada et al., 1998 ). Shallow-marine glycymerid beds areH. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 931
intercalated in the basal part of the Oku-utsunai Formation,
but the author did not find any good index megafossils.Judging from the microfossil biostratigraphy and theconformable transgressive trend from the Oku-utsunai toUtsunaigawa Formations, the uppermost Maastrichtianabove the ammonite extinction horizon and the Danian aswell as the K/T boundary seem to have been eroded away atthis unconformity. Its time gap is estimated as 7–8 m.y.,according to a time scale by Gradstein et al. (1995) .
6. Discussion: sedimentary history of the Yezo
forearc basin
Forearc basin sedimentation is closely related to plate
tectonic movements, arc volcanism and forearc accretionaryprocesses on a large scale. Global eustasy and local relativesea-level changes, sediment supply, local tectonics such as
basin subsidence and uplift are also important controlling
factors, though it is often difficult to discriminate theireffects on each other. The temporal and spatial distributionof sediments and their facies successions in Northeast Japanto Sakhalin indicate the sedimentary history of the Yezoforearc basin during Albian to Paleocene time.
Stratigraphic sections in Hokkaido are generally a few
times thicker than those in North Honshu, though totalthickness varies area by area. For instance, column 11(Oyubari) in Hokkaido reaches 5000 m in maximumthickness. All continuous sections in Hokkaido exceed afew thousands meters in thickness. On the other hand, themaximum is about 2700 m for (columns 22 (off Iwaki) and23 (Joban) in north Honshu), though it is a composite valuebecause of a lack of continuous sections. Every section in
north Honshu is stratigraphically short ranging in compari-
son with Hokkaido. This is partly because of a post-Paleocene unconformity and subaerial erosion caused bytectonic movements related to subduction processes(Kimura, 1997 ). Much thicker sediments in Hokkaido
suggest that the subbasin also subsided a few times fasterthan in the Kitakami and Joban subbasins. Very thick andlong-ranging sedimentary successions including shallow-marine to non-marine strata such as the Mikasa Formationand the Hakobuchi Group clearly indicate that the basinevolution was tectonically controlled by a fast rate ofrelative plate motion ( Engebretson et al., 1985 ) as well as
local tectonics. They may be related to rapid growth of theeastward accretionary complexes represented by the Idon-nappu and Hidaka Belts (e.g. Kimura, 1997; Ueda et al.,
2000 ). On the contrary, the growth rate may had been lower
in north Honshu than in Hokkaido, though contempora-
neous accretionary sediments have not been found beneaththe continental slope off the Pacific Ocean ( Kimura, 1997 ).
In the first second-order upward-shallowing cycle, thick
terrigeneous clastics were deposited as turbidites of theLower and Middle Yezo Groups in Hokkaido. Shallow-marine to paralic facies were developed only in the westernmarginal part of the basin during late Albian to Turonian,
especially the Mikasa Formation in central Hokkaido. TheMikasa is composed of three third-order depositionalsequences, each of which further include three fourth-order sequences. These third- and fourth-order sequencesshowing upward-coarsening facies successions probablyreflect repetitive delta progradation ( Ando, 1990b, 1997) .
Furthermore, the southward shift of the delta system withinthe distribution area of the Mikasa is documented by itsstratigraphic range ( Ando, 1990a,b, 1997 ;Fig. 3 , columns
8–10). The stratigraphic range changes from the uppermostAlbian-Cenomanian in the northern Ashibetsu area (column8), to the Turonian in the southern Yubari area (column 10).This is possibly related to local tectonics in the westerncentral margin of the basin.
During the Coniacian to Santonian, the Hokkaido
subbasin was dominated by offshore marine muddysedimentary environments (outer shelf to upper continentalslope) of the Upper Yezo Group. This transgressive andhighstand stage has been called the Urakawan transgression(Yabe, 1926 ). This means that the basin was generally
starved of terrigeneous clastics, though turbidites anddebris-flow deposits such as the Tsukimi Formation arelocally developed in the Ashibetsu area ( Tanaka, 1963 ;Fig.
3, column 8). However, the transgressive and highstand
trends in the Kitakami and Joban subbasins are slight for theConiacian to Santonian. Some differentiation within the
Yezo basin may have taken place.
The predominance of shallow-marine to fluvial plain
sediments represented by the Hakobucih Group inHokkaido and its correlatives in north Honshu indicatesthat a large amount of coarse terrigeneous clastics had beensupplied into the Yezo basin during Campanian andMaastrichtian and almost filled it up until the lateMaastrichtian. This change of sedimentary setting isthought to have been caused by uplift of the continentalmagmatic arc probably relating to plate motion ( Ando,
1997 ). According to Engebretson et al. (1985) , the azimuth
change occurred from NNE to W or WNW for the Izanagiand Kula plates about 85 Ma ( Fig. 3 ). They may be also
related to rapid growth of the accretionary complexes inthe Idon-nappu and Hidaka Belts (e.g. Kimura, 1997; Ueda
et al., 2000 ;Fig. 3 ).
The uppermost Maastrichtian to Lower Paleocene,
including the K/T boundary, seem to be absent throughoutthe basin, judging from the biostratigraphy on dinoflagel-lates, nannofossils, ammonites and inoceramids. There wassubaerial erosion during the K/T gap, though no heavytectonic disturbance. After the K/T gap, marine to non-marine Late Paleocene sediments were sporadically devel-oped in the Nakatonbetsu and Yubari areas in Hokkaido,and part of the offshore Kitakami and Joban subbasins.Because lithological changes at the gap are not detected, thesame basin framework would have been kept to someextent. Furthermore, the lithological similarity between theuppermost Hakobuchi Group and the Haboro Formation
H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 932
suggests that the latter is an element of the Yezo basin. In
other words, the uppermost part of the Hakobuchi Group,the Paleocene strata recognized in offshore drilling off northHonshu and the Haboro Formation all represent the finalfilling phase of the Yezo basin. After the late Eocene,basinal tectonic setting as well as sedimentary facies inHokkaido changed, as pointed out by Iijima (1996) and
Kurita and Yokoi (2000) .
Imaoka et al. (1999) reconstructed the Cretaceous to
Tertiary igneous activity in the Honshu Arc by using Rb–Srwhole rock isochron ages. In north Honshu, igneous activityoccurred during 135–100 Ma, judging from widely-dis-tributed granite and narrowly-associated andesitic volcanicsin the Abukuma, South and North Kitakami, and OshimaBelts ( Fig. 3 ). Because of the lack of radiometric age data
for 89–76 Ma in the four belts, the activity seems to haveceased in north Honshu, though not in Southwest Japan.Again it took place in north Honshu from 75 to 71 Ma, andcontinued until the Paleogene though not as extensive as inthe Early Cretaceous. The predominance of arkosicsandstone in the Futaba Group means that the LowerCretaceous granite batholith was already exposed on thehinterland in the Coniacian. Acidic tuff layers suddenlyincreased in the Campanian strata throughout the basin. Thisappears to reflect the latest Cretaceous to Paleogene igneousactivity. On the other hand, the Mikasa Formation has amoderate amount of sand and gravel derived from
intermediate to basic igneous rocks, reflecting the lower
Cretaceous basaltic to andesitic volcanic rocks distributed inthe westerly Rebun-Kabato Belt ( Okada and Matsumoto,
1971; Nagata et al., 1986 ).
Ito and Masuda (1992) briefly pointed out the synchro-
nicity of the spatial-temporal sediment distribution inJapanese basins. They thought that basin evolution andmajor episodes are attributed to the relative motion of oceanplates along the Japanese convergent margin, especially inthe case of the Upper Cretaceous, Izanagi and Kula plates(Engebretson et al., 1985 ). As to the second-order cyclicity
of the Yezo forearc basin-fill, this synchronism can bedemonstrated as discussed above. However, the third-orderdepositional sequences which are well developed inshallow-marine to non-marine facies cannot be explainedby such large-scale synchronism. The stacking patterns ofthird-order sequences and the timing of sequence bound-aries and key surfaces bounding systems tracts appearconsiderably concordant with the oscillation patterns of theHaq curves ( Haq et al., 1988 ), though some inconsistencies
can be recognized for the Maastrichtian in the upper part ofthe Hakobuchi Group. As discussed in Ando (1997) , the
average sedimentation rate for four strata (Hakobuchi, Kuji,Futaba Groups and Mikasa Formation) is broadly similardespite their different tectonic settings. These data suggestthat the third-order sedimentation may have been controlledby not only local tectonic movement but also globaleustasy even in this active-margin setting. However, it isdifficult to identify the main controlling factor amongseveral sedimentary constraints for the fourth-order
sequences or parasequences, which are developed in HST
of some third-order sequences. We need further information
and consideration for the better understanding of thesedimentary history of the Yezo forearc basin and the
geologic factors controlling the depositional sequences.
7. Conclusions
The temporal and spatial distributions of the sediments
and their facies successions in Northeast Japan to southSakhalin during Albian to Paleocene time indicate that two
second-order upward-shallowing cycles equivalent to
megasequences are developed throughout the Yezo forearcbasin. These cycles are well developed in Hokkaido but are
obscure in north Honshu. Each cycle includes a few third-
order and several fourth-order depositional sequences.Shallow-marine to paralic facies in the late stage of the
first cycle are represented by the Mikasa Formation (late
Albian to Turonian), while the eustatic curves ( Haq et al.,
1988 ) reach the highstand maximum during the late
Cenomanian. This discrepancy may suggest local tectonic
movement only in the western margin of western centralHokkaido. The Campanian to Maastrichtian shallow-marine
to non-marine facies of the second cycle are characterized
by the Hakobuchi Group in Hokkaido and its correlatives innorth Honshu. This large-scale shallowing trend may reflect
the directional change of the Izanagi and Kula plate motions
in the early Campanian. Because the uppermost Maastrich-tian to Upper Paleocene, including the K/T boundary, are
possibly absent everywhere in the Yezo basin, tectonic
uplift causing subaerial erosion seems to have occurred. Thethird-order sequences are thought to have been controlled by
global eustasy as well as local tectonics, since the stacking
patterns of depositional sequences and the timing ofsequence boundaries and key surfaces appear considerably
concordant with the oscillation patterns of the Haq curves
even in this active-margin setting.
Acknowledgements
The author thanks JNOC (Japan National Oil Corpor-
ation) for the permission to use drill hole data in this work. I
acknowledge the following persons: Tadashi Oguro, T.(JNOC), Takao Iwata (Teikoku Oil Co. Ltd.) and Hiroshi
Kurita (Niigata University) for providing much information
of the exploration data, Andrew Martin of Japan NuclearCycle Development Institute for improving the manuscript.
Special thanks go to William S. Elliott, Jr. and Bruce Hart
for helpful reviews. This research has been supported in partby Grant-in-Aid for Scientific Research from the Ministry of
Education, Science and Culture of Japan (No. 07640619,
10640446).H. Ando / Journal of Asian Earth Sciences 21 (2003) 921–935 933
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Memoirs of the Faculty of Science, Kochi University Series E, Geology. Vol. 8. p. 1-70, pls. 1-6, March, 1987. <0.01 IGNEOUS AND METAMORPHIC PETROLOGY OF THE HOROKANAI OPHIOLITE IN THE KAMUIKOTAN ZONE. <0.01 0.221 <0.01 ly emplaced as a nappe onto the Kamuikotan subduction complex in Hokkaido, Japan, during a late Mesozoic time. The ophiolite is superposed in the following general sequence, starting from the top and working down; composed of basaltic pillow lava, hyaloclastite, tuf, lava flow, banded amphibolite and massive amphibolite, and ultramafic rocks (2000-2500 m thick) consisting of pyroxenite, dunite and harzburgite with minor lenses or layers of gabbroic rocks. Of these, the banded and massive amphibolites were derived from basaltic hyaloclastite-massive lava flow and gabbroic rocks, respectively. The dikes of dolerite intruded the mafic rocks as feeder-channels for the extrusive basaltic rocks, but the typical sheeted dike complex is absent, and also the dikes of plagiogranite occur near the boundary between the banded and massive amphibolites (i.e. between the basaltic and gabbroic rocks). Furthermore, near the boundary between the massive amphibolite and the ultramafic rocks occur layered cumulate rocks characterized by the alternation of dunite, wehrlite, clinopyroxenite, oliving-gabbro and gabbro. Geochemically, the basaltic rocks have relatively low FeO*/MgO values and high Cr and Ni contents, and show a tholeitic fractionation trend defined by possitive correlations of V and TiO2 with FeO*/MgO, while the gabbroic rocks have lower FeO*/MgO values and higher Cr and Ni contents than the basaltic rocks, and are de- pleted in “incompatible" elements (Y and Zr), as a result of their primitive nature. The tholeitic nature of the basaltic rocks is also supported by the relict pyroxene chemistry. The plagiogranites have a distinct geochemistry rs s r m o Ni contents. The abundance of some minor elements as well as the relict spinel chemistry suggest that the Horoka- nai ophiolite has a close affinity with present-day abyssal tholeites. Although the top sedimentary rocks exhibit no metamorphic nature, the underlying mafic rocks display signi- mineral zones are distinguished in the ascending order of metamorphic grade; Zone A (the zeolite facies): zeolites + chlorite + albite + pumpellyite, Zone B (the greenschist facies): albite + chlorite + actinolite + epidote, Zone C (the greenschist-amphibolite transition facies): albite + oligoclase + actinolite + hornblende + chlorite + epi- dote, and Zone D (the amphibolite to partly granulite facies): hornblende + calcic plagioclase + clinopyroxene + orthopyroxene. These four mineral zones are areally mapped, indicating that the grade of metamorphism increases down the ophiolite sequence. The basal ultramafic rocks commonly retain igneous texture, but their mineral com- 1,914 Coexisting metamorphic minerals from several tens of samples covering all the mineral zones were analysed by means of an electron-microprobe analyser, and the following compositional variations with metamorphic grade are observed; (1) the AlzO; contents of pumpellyite and epidote, and the Mg/(Mg+Fe*) ratio of chlorite increase Zone C albite and oligoclase coexist defining the peristerite gap with the shape of a two-phase binary loop, (3) the Al'V content of amphibole increases from Zone B to Zone D, but in the medium-grade part of Zone C actinolite and hornblende coexist showing the compositional gap attributed to a solvus, and (4) the CaO content of clinopy- a p xoo n n perimental studies and using thermometers of two-pyroxenes and olivine-spinel, the metamorphic facies series of the Horokanai ophiolite corresponds to the low-pressure type with a temperature range of 100-750°℃. Sujiyu 2 HIDEO ISHIZUKA The metamorphic nature of the Horokanai ophiolite is broadly comparable to that inferred for ocean-floor o ie s u soi s a jo sd or si u r its scarceness in ocean-floor metamorphism. The metamorphism of the Horokanai ophiolite may have taken place near or beneath oceanic spreading ridge where a steady-state magma chamber exists to generate an ophiolite, and has a very slow cooling rate and may be recrystallized. This is a kind of autometamorphism. Also, a hydrothermal o p n jo uo r n e d q u jo m recrystallization may depend upon the spreading rate of the oceanic ridge, and the Horokanai ophiolite may have s n o q e 1. INTRODUCTION that postdated the primary igneous origin, and The mafic-ultramafic rock association with a cap- there are two types of metamorphic facies series: ping of minor sediments has been known to occur (1) the low-pressure type as recognized in in orogenic belts in the world. A term “ophiolite" ophiolites of the Mediterranean (Gass & Smew- is re-defined to denote such an association (Pen- ing, 1973; Spooner, 1974; Spooner & Fyfe, 1973; rose Conference, 1972). In a completely de- Spooner et al., 1974; Pamic et al., 1973), Alps veloped ophiolite, the rock-types occur in the fol- (Mevel et al., 1978), Oman (Coleman, 1977), lowing sequence, starting from the top and work- Newfoundland (Coish, 1977), Chile (DeWit & ing down: (1) Sedimentary rocks typically includ- Stern, 1976; Stern et al., 1976; Elthon & Stern, ing radiolarian cherts, thin shale interbeds and 1978; Stern & Elthon, 1979), Taiwan (Liou, 1979; minor limestones; (2) Mafic rocks composed of (a) Liou & Ernst, 1979), Borneo (Hutchison, 1975, basaltic rocks commonly pillowed,(b) doleritic 1978), California Cast Range (Evarts & Schiff- dikes or sheets, and (c) gabbroic rocks with or man, 1983), and Hidaka Western Zone (Miyashita without cumulus pyroxenites and peridotites; (3) et al.,1980; Miyashita, 1981,1983), and (2) the Ultramafic rocks consisting of variable proportions high-pressure type as recognized in ophiolites of of harzburgite, lherzolite and dunite (more or less Franciscan (Page, 1981), New Caledonia (Black & serpentinized). Brothers, 1977), Alps (Ernst, 1973), and Sanbaga- The origin of ophiolites had been long con- wa (Miyashiro, 1973b). Furthermore, it has been troversial (for historical review, see Coleman, also recognized that ophiolites of Alps (Dietrich et 1977), but in the early 1970s the concept of the al., 1974) and Kurosegawa (Maruyama et al., plate tectonic theory enabled to interpret 1978) have textural and mineralogical evidences ophiolites as tectonically y emplaced portions of that indicate polymetamorphism characterized by ancient oceanic materials generated by the same the early low-pressure type and the late high- mechanism presently governing the formation of pressure type. The medium-pressure type of the new oceanic materials along spreading oceanic ophiolite metamorphism is very rare, but has been ridges (e.g. Moores & Vine, 1971; Coleman, recently reported from the Yakuno ophiolite (Ishi- 1971). However, based on petrochemistry, con- watari, 1985). siderable controversy has again appeared concern- The metamorphosed ophiolites of the high- ing the tectonic setting in which ophiolites can be pressure type are closely associated with high- generated (Jakes & Gill, 1970: White et al., 1971; pressure metamorphic terrains, and more com- Ewart & Bryan, 1972; Miyashiro, 1973a), and monly occur as dismembered ophiolites within Miyashiro (1975) emphasized that ophiolites can melanges. On the basis of the current scheme that be formed in tectonic setting other than that of the high-pressure metamorphism has evolved spreading oceanic ridges, such as island arc or along a convergent plate boundary (Coleman, oceanic island. As a consequence, a number of 1971; Ernst, 1974), the ophiolite metamorphism of ophiolites have been reported to be of island arc the high-pressure type could be closely related to origin (Phelps & Ave Lallement ,1980; Menzies the tectonic history of ophiolite emplacement into et al., 1980; Gerlach et al., 1981). Furthermore, the orogens of continental margins. Jakes & Miyake (1984) recently suggested that a The metamorphosed ophiolites of the low- forearc region is another candidate for the forma- pressure type are generally unrelated to the high- tion of an ophiolite. pressure metamorphic terrains, but some of them, On the other hand, ophiolites have been most notably that in the California Coast Range, more or less modified by metamorphic processes occur as nappes in the high-pressure metamorphic Igneous and metamorphic petrology of the Horokanai ophiolite terrains. The relatively intact suite of the ophiolite of the primary (igneous) texture and mineralogy stratigraphy has been commonly preserved in the by metamorphism makes it somewhat difficult to metamorphosed ophiolites of the low-pressure study this problem. type. In contrast with the high-pressure meta- This paper is part of my doctoral thesis sub- morphism, there are continuing uncertainties on mitted to the Kyoto University, of which some the processes responsible for the low-pressure have been published (Ishizuka, 1980a, 1980b, ophiolite metamorphism. However, the fact that 1981a, 1981b,1985; Ishizuka & Imaizumi,1980; this type of metamorphism modifies only the Ishizuka et al.,1981,1983a, 1983b, 1984; Banno ophiolite mineral assemblages but not neighbour- & Ishizuka, 1986). ing country rocks suggests that its metamorphic stage predated the ophiolite emplacement onto the present portion. In this connection, it may be 2. OUTLINE OF GEOLOGY noted that recent progress on petrology of drilled 2-1. Geological Setting or dredged metamorphic rocks from ocean-floor In the axial zone of Hokkaido, the Kamuikotan has revealed that much of the oceanic crust has terrain extends from south to north for a length of suffered metamorphism (ocean-floor metamorph- ‘ss 1o g 0 ism) (Melson & van Andel, 1966; Cann & Fun- further it extends northwards to Sakhalin (Fig. nell, 1967; Cann, 1969; Ploshko et al., 1970; 1A). In the current scheme of metamorphic belts Miyashiro et al., 1971; Aumento et al.，1971; of the Japanese islands, this terrain is regarded as Bonatti et al., 1975; Helmstaedt, 1977; Mevel, a high-pressure metamorphic belt of the jadeite- 1981; Ito & Anderson, 1983; Honnorez et al., glaucophane type (Miyashiro, 1961; Hashimoto et 1984). This recrystallization includes the metamor- al., 1970), commonly associated with a great phic facies series of the low-pressure type. The na- quantity of ultramafic rocks. In addition to ture of the low-pressure ophiolite metamorphism glaucophane, lawsonite and jadeite + quartz, all is broadly similar to that inferred for ocean-floor known at those times, aragonite has been found to metamorphism, and it has been widely accepted be a widespread metamorphic mineral coexisting that ocean-floor metamorphism is the probable with lawsonite (Gouchi & Banno, 1974; Shibakusa mechanism for the low-pressure ophiolite meta- et al., 1977). Metamorphic zonal mapping has morphism. However, the substance of ocean-floor been attempted by Banno & Hatano (1963), Taza- metamorphism is not clear, because there is no ki (1964), Shibakusa (1974), Herve (1975), Aguir- direct knowledge on the structural and strati- re (1977), Nakano (1981) and Gouchi (1983), all graphical relationships among the dredged meta- authors concluding that these rocks belong to the morphic rocks from the ocean-floor; the deepest high-pressure facies series, but the details of its complete reference section of oceanic crust is of metamorphic geology has been still in dispute. Re- 1075.5 m which has been established by Legs 69, cently, Imaizumi (1983) and Maekawa (1983) have 70 and 83 of the Deep Sea Drilling Project revealed that the majority of the Kamuikotan (Anderson et al., 1982). metamorphic rocks are of a melange formed by The Horokanai ophiolite studied here occurs accretion, subduction and related processes (e.g. as a nappe emplaced onto the Kamuikotan high- sedimentary recycling associated with a trench- pressure metamorphic rocks in Hokkaido, Japan. olistostrome, and tectonic transportation carried It has a geological feature in common with other by serpentinite). Mega-fossils (brachiopods, corals ophiolite nappes, and most importantly it has a and stromatoporoids) of late Mesozoic age (Hashi- complete metamorphic gradation within the low- moto, 1971; Matsumoto & Okada, 1971) and pressure facies series, and hence it may provide an micro-fossils (radiolarians) of Valanginian to Hau- excellent opportunity for understanding the terivian age (Nakaseko, 1979; Okada et al., 1982; mechanism of the low-pressure ophiolite meta- Watanabe, 1982) have been discovered from the morphism. The present thesis is mainly focused on Kamuikotan metamorphic rocks, while K-Ar ages the metamorphic petrology of the Horokanai of the Kamuikotan metamorphic rocks have been ophiolite, along with the discussion on the origin determined to be 109-120 Ma (Bikerman et al., of its metamorphism. The problem of the igneous 1971) and 72-145 Ma (Imaizumi & Ueda, 1981). origin of the Horokanai ophiolite is also discussed Therefore, the Kamuikotan metamorphic rocks with using bulk rock chemistry and relict mineral represent a subduction complex evolved during a chemistry, even though the general modification late Mesozoic time. HIDEO ISHIZUKA (A) 145°E (B) NoOS 140°E Sakhalin NoS7 Hokkaidd Kamuikotan Zone No0t (C) 0.5- 0 (km) EXPLANATION Neogene Volcanic Rocks Cretaceous Yezo Group Ophiolitic Mafic Rocks & Related Sediments Uitramafic Rocks Kamuikotan Subduction Complex H:HOROKANAI K:KAMUIKOTAN GORGE A:ASAHIKAWA (=ASAHIGAWA) Fig. 1. A: Regional setting of the Kamuikotan zone in Hokkaido and northwards to Sakhalin, B: Geological outline of the Kamuikotan zone in the Kamuikotan gorge to Horokanai area, and C: The cross- section along A-B-C of (B). Igneous andmetamorphic petrology of the Horokanai ophiolite 5 On the other hand, it has long been known The sequence of the Horokanai ophiolite is that another group of metamorphic rocks (mainly faulted and folded to hinder a quantitative esti- amphibolites) also occurs associated with ultrama- mate of its original thickness. However, the gener- fic rocks in the same terrain (Hunahashi, 1944; al sequence of the ophiolite is confirmed by the Suzuki & Suzuki, 1958; Igi et al., 1958; Igi, 1959; detailed field survey using 1/5000 topographic Banno & Hatano,1963; Watanabe,1965), and maps, and it is schematically shown in a columnar that these rocks contain hornblendes with low- section of Fig. 2. The ophiolite is superposed in pressure compositions (Haramura, 1963), but their the following general sequence, starting from the significance has been largely ignored. Recently, top and working down; sedimentary rocks (30-50 Banno et al. (1978), Asahina & Komatsu (1979) m thick), mafic rocks (2000-2300 m thick) with and the present author have re-examined these dikes of dolerite and plagiogranite, and ultramafic amphibolites and adjacent rocks, and discovered rocks (2000-2500 m thick); the modes of occurr- that there is a complete metamorphic gradation ence of these rocks are briefly described below. within the low-pressure facies series between the amphibolites and overlying metabasites. It follows 2-2-1.Sedimentary Rocks that these low-pressure metamorphic rocks actual- The sedimentary rocks comprise mainly cherts, ly form an integral part of the Kamuikotan ter- and they can be divided into three types: (1) mas- rain. sive chert, (2) bedded chert, and (3) volcanic These facts allow the interpretation of the fragment-bearing chert. The type (1) chert is vari- Kamuikotan terrain as a zone where the high- able in color, although it is commonly greenish to pressure and low-pressure metamorphic rocks pale greenish but in some cases greyish or reddish, have been tectonically mixed up, and it is better and usually exhibit dim glassy luster. The type (2) not to refer to this zone as the Kamuikotan ter- chert, which is the most common type, is greenish rain in order to avoid the impression that it is a to reddish in color, and alternates with greenish to single regional high-pressure metamorphic belt in reddish shale to form bedded structure (Plate 1A). the ordinary sense. Therefore, it is proposed that The type (3) chert is greenish in color, and irre- this terrain be referred to simply as the Kamuiko- gularly includes many spheroidal to subspheroidal tan zone. volcanic fragments derived from pillow basalts. The Horokanai area studied in this thesis is Stratigraphically, the type (1) and (2) cherts over- located in the central part of the Kamuikotan lie the basaltic pillow sequence, while the type (3) zone, about 30 km northwest of Asahikawa chert is commonly intercalated with pillow basalts. (=Asahigawa), and is important as one of the Fossil radiolarians available on age deter- most convincing and informative occurrences of mination are abundant in the type (1) massive both the high-pressure and low-pressure metamor- chert (Plate 1B). Under the scanning electron- phic rocks within the Kamuikotan zone; the re- microscope (SEM) and light microscope (LM), a gional geological sketch map is shown in Fig. 1B total of eight genera and four species of radiola- and the cross section in Fig. 1C. Of these, the rians were identified by Dr.M.Okamura. Among low-pressure metamorphic rocks and related ultra- the species,F Parvicingula hsui has characteristic mafic rocks and chert construct an ophiolite sequ- pore pattern of postabdominal chambers, which is ence in the sense of Penrose Conference (1972), common in Zone 2B to Zone 3 (early Tithonian) and hence they are called “Horokanai ophiolite" from the chert of the Coast Range Ophiolite, hereafter. California (Pessagno, 1977). According to Baum- garter et al. (1980), Mirifusus mediodilatatus is a 2-2. Field Observations junior synonym of M.baileyi Pessagno (1977), The geological map of the area studied here in de- which is widely recognized in Upper Jurassic stra- tail is shown in Fig. 2. The area is underlain by ta. Praeconocaryommamagnimamma is also the Kamuikotan metamorphic rocks, the Horoka- known to range from late Jurassic, but is assumed nai ophiolite, the Cretaceous Yezo Group (un- to have its maximum development during Zone metamorphosed sediments), and the Tertiary vol- 2A (early Tithonian) by Pessagno (1977). Tricolo- canic rocks. The Horokanai ophiolite is in fault- capsa sp. and Zhamoidellum sp. are commonly contact with the Kamuikotan metamorphic rocks found in the late Jurassic chert from the Sanbosan and the Yezo Group, and is locally covered by the Group in Shikoku (Aida, 1982). Accordingly, it is Tertiary volcanic rocks. suggested that the chert of the Horokanai HIDEO ISHIZUKA ophiolite is correlated to late Jurassic, most prob- underlain by massive amphibolite. This amphibo- ably early Tithonian in age. lite sequence may represent a complex of basaltic hyaloclastite-lava followed by gabbro. In the 2-2-2. Mafic Rocks above sequence, numerous dikes of dolerite occur, The mafic rocks consist of basaltic lavas and a but the typical sheeted dike complex described complex of basaltic hyaloclastite-tuff-massive lava elsewhere (e.g. Coleman, 1977) is absent.Near the flow. Deeper in the sequence, the nature of the boundary between the basaltic and gabbroic rocks, precursors is obliterated by metamorphic recrystal- i.e. between the banded and massive amphibo- lization, but we found that banded amphibolite is lites, dikes of plagiogranite occur. These dikes of Zone ANATION Vvy]NeogeneVolcanicRocks YezoGrou ? nuikotanMetamorphie Rocks HOROKANAI OPHIOLITE RadiolarianChert 品 mas Cumulate Rocks Uitran dunite foliated harzburgite harzburgite entinite) Mixed-Up Zone) (barbsonupperplate 3(km) ini Fault Lithological Fig. 2. Geological map of the Horokanai ophiolite (after Ishizuka, 1980a, 1985). The thin arrow in the inset points to the locality of the Horokanai ophiolite. Igneous and metamorphic petrology of the Horokanai ophiolite / dolerite and plagiogranite are also metamorphosed this type of hyaloclastites may have been formed to the same grade as the adjacent host metaba- by collapse of pillow surfaces due to rapid cooling sites. during the formation of the pillow pile (Rittmann, The pillow lavas commonly occur as close- 1962). In the hyaloclastites, the basaltic tuffs occur packed type with cogenetic interpillow materials, as several layers, ranging from 1 to 10 m in thick- and their outcrops appear bulky, rounded and ness (Plate 1G). They are pale green to pale gray hummocky (Plate 1C). Individual pillows are in color, but commonly fragile due to weathering. generally ellipsoidal in cross section with the max- The contact between the tuffs and hyaloclastites is imum and minimum diameters ranging from 30 to irregular and sometimes unclear. Lithologically, 80 cm and from 20 to 50 cm, respectively. Occa- the tuffs resemble the hyaloclastite matrix, and sionally, large bolster pillows of about 5 m long their deposition may have occurred intermittently and 2 m wide show flattened and tabular form, during the period when the episodic submarine and some are elongated or sausage-like pillows in sliding of pillow piles took place to form the the sense of Sigvaldason (1968). A glassy rind of hyaloclastites. Also, the basaltic massive lavas several millimeters in thickness develops at the occur as a few layers intercalated within the pillow margin (i.e. chilled margin). On the pillow hyaloclastites, but the contact between the mas- surface, the glassy part is traversed by a polygonal sive lavas and hyaloclastites is obscure (Plate 1H). network of shrinkage cracks (i.e. chilled cracks), They appear blocky on the weathered surfaces, which extends inwards forming radial columnar and sometimes exhibit flow layering. joints. The visible vesicles or amygdules are rare The banded amphibolite is characterized by in the glassy pillow margin, but present sporadi- banded structure composed of leucocratic cally in the crystalline pillow core. The pile of pil- (plagioclase-rich) and melanocratic (amphibole- lows generally lacks other than crude stratifica- rich) bands (Plate 2A). Each band is subparallel tion, and consequently it is very difficult to obtain and ranges from 5-10 mm in thickness. However, the precise attitude of the pillow sequence in the such banded structure tends to be weak down- field. However, a tendency for pillows to flatten wards, and the rocks appear to be massive, and in the place of deposition results in an approxi- hence the contact between the banded amphibo- mate "bedding". In this respect, it is also notewor- lite and underlying massive amphibolite is tran- thy that some of pillows are balloon-shaped with sitional. The massive amphibolite commonly circular upper surface and prominent tails plasti- shows gneissose structure (Plate 2B). cally protruding downwards, fitting into re-entrant The dikes of dolerite, ranging from 0.5 to 5.0 spaces of underlying pillow surfaces (Plate 1D). m in thickness, occur with sharp contact against pue spiemdn xaauos) .sdon paains-qoous, yons host rocks (Plate 2C). The chilled margins and tails downwards) provide a criterion to differenti- columnar joints develop well, and the grain size ate the top and bottom of the pillow sequence. tends to grow coarser from the margin to the core The hyaloclastite is a tuff breccia-like rock of the dike. The dikes of dolerite identified in the containing many volcanic fragments embedded in study field are 37 in number. Also, dikes of plag- cogenetic tuffaceous matrix, and usually displays iogranite, ranging from 10 to 50 cm in thickness, weak metamorphic foliation (Plate 1E). The vol- develop with sharp contact against the host rocks, canic fragments, ranging from angular to subangu- but the chilled margin is unclear (Plate 2D). Occa- lar or subrounded in shape and from 1 cm or less sionally the multiple dikes of dolerite are observed to more than 10 cm in size, are largely derived (Fig. 3A), probably formed as follows; one dike from pillow basalts with minor doleritic rocks, of has intruded, forming the chilled margin against -l Asi e Aq pa Aed sne sos ym the host rock, and then next dike has intruded in led) rind. No sedimentary structure such as bed- the inner part of the previous one and in turn ding, lamination and grading structures is formed the chilled margin against the host, pre- observed, suggesting that this type of hyaloclas- vious dike rock. Very rarely it is observed that the tites is an in situ hyaloclastite formed by breaking dike intruding the pillow pile passes to pillow of the pillow pile after its complete consolidation, lavas in its margin (Fig. 3B), suggesting that most but not a re-worked hyaloclastite caused by large- of dolerite dikes are feeder-channels for the extru- scaled turbidity current (Kawachi et al., 1976). sive basaltic rocks. Occasionally, numerous fine-grained glassy pillow fragments are observed (Plate 1F), indicating that 2-2-3. Ultramafic Rocks 8 HIDEOISHIZUKA (A) WSw NE （B) ：orthopyroxenite : dunite SW >NE 1m : harzburgite : gabbro Fig. 3. A: A sketch showing multiple dolerite dikes in hyaloclastite, in which the numbers (1, 2 :pegmatitic and 3) represent the order of intrusion, and hornblende-gabbro s$s]: serpentinite B: A sketch showing transition of dolerite dike to pillow lavas (after Ishizuka, 1980a). K.M.R:Kam The ultramafic rocks are largely divided into the Fig. 4. Lithological map of the southern ultramafic northern and southern bodies in relation to dis- rocks (after Ishizuka, 1980a). tribution; the condition of outcrops is considerably better in the southern body than the northern one. As a whole,they underlie the massive Generally, the ultramafic rocks are serpenti- amphibolites, but near the boundary between the nized to some degree, but their parental peridotite ultramafic rocks and massive amphibolites occur types are inferred from relict texture and mineral- layered rocks as many boulders on the river bed ogy to have been dunite and harzburgite with (Plate 2E). These layered rocks consist mainly of minor orthopyroxenite; the distribution of these olivine, clinopyroxene and plagioclase with minor rock-types is mapped in the southern body, as orthopyroxene, of which the olivine/clinopyroxene illustrated in Fig. 4. Stratigraphically, they com- ratio is highly variable to give rise to variation of prise orthopyroxenite (50 m thick) at the top fol- rock-types such as dunite, wehrlite, clinopyroxe- lowed by dunite (500 m thick) and then harzbur- nite, olivine-gabbro and gabbro (Plate 2F). Each gite (2000 m thick), of which the basal harzburgite of these rock-types forms a rhythmically alternat- is further divided into upper foliated and lower ing and subparallel layer, ranging in thickness massive types; the younging direction of the ultra- from a few centimeters to less than 1.0 m. Mineral mafic rocks are mainly determined by means of grading of olivine or plagioclase is sometimes dis- grading structure definedby olivine and/or tinct. These( observations indicate the layered orthopyroxene crystals. The boundary between rocks to be cumulates. The filed relationships these rock-types is characterized by layering struc- among these rock-types are, however, obliterated ture(Plate 2G); i.e. the boundary between by poor exposure, even thoughAsahina orthopyroxenite and dunite is marked by gradual Komatsu (1979) reported that the dunite and decrease of orthopyroxenite layer and increase of wehrlite predominate in the lower part, the cli- dunite layer, and that between dunite and harz- nopyroxenite in the middle part, and the gabbro burgite by gradual decrease of dunite layer and in- in the upper part of the cumulate sequence. crease of harzburgite layer. No clear "unconformi- Igneous and metamorphic petrology of the Horokanai ophiolite ty” is present between these rock-types. On the occurred well after the cessation of the two dis- other hand, the seams rich in Cr-spinel usually de- tinct types of metamorphism as evidenced by the velop in the dunite, and also the layers of olivine- following reasons. First, the contact between the gabbro and gabbro are often accompanied by the ophiolite and the Kamuikotan metamorphic rocks basal harzburgite, but the continuity of these is a thrust fault where the sheared serpentinites seams or layers is poor (Plate 2H) derived from the basal harzburgite of the ophiolite occur including the fragments of the Kamuikotan 2-3. Structural Features and Tectonic Emplace- metamorphic rocks as fault breccias. The over- ment lying ophiolite has the top chert of Tithonian age The dip-strike C directions of the Horokanai (about 145 Ma) and the metamorphic hornblendes ophiolite are obtainedbythe bedded plane of of 176-186 Ma (Ar40-Ar39 ages: Takigami, 1983 type (2) bedded chert, the foliation plane of the personal communication) while the underlying hyaloclastite, the banded plane of the banded Kamuikotan metamorphic rocks have the meta- amphibolite, and the layering plane of the ultra- morphic ages of 72-145 Ma (K-Ar ages for musco- mafic rocks; these planes are subparallel to each vites: Imaizumi & Ueda, 1981). Second, the press- other. They strike N10°W-N30°E and dip 40°E- ures and temperatures of mineral equilibrium are 60°E (Fig. 5). On the other hand, the Horokanai discontinuous between the ophiolite and the ophiolite occurring in the northwestern part of the Kamuikotan metamorphic rocks. Fig. 6 illustrates Horokanai area also strikes N10°w-N30°E but the process of the tectonic emplacement of the dips 40°W-60°w. It follows that the Horokanai ophiolite onto the Kamuikotan metamorphic rocks ophiolite is distributed symmetrically with the as a nappe, in which the “Mixed-Up Zone" is also Uryu-River as an axis (Figs. 1B and 1C). shown to be formed during the ophiolite emplace- As shown in Fig. 2, the Horokanai ophiolite ment. is divided into the northern and southern masses These structural, lithological and age rela- by “Mixed-Up Zone". It is a low-angle fault zone, tionships are similar in some aspects to those in extending northwest to southeast with about 250 western California where the Coast Range m width and dipping 30-40° northeast, in which Ophiolite (low-pressure meta-ophiolite) has been numerous angular to subangular blocks were thrust onto the Franciscan complex (high-pressure mixed up; the blocks range in size from 50 cm or subduction complex) (e.g. Page, 1981). In Califor- less to more than 10 m, and comprise various nia the associated forearc-basin sediments with the rock-types of the ophiolite, but do not include Franciscan subduction is considered to be the blocks derived from the Kamuikotan metamorphic Great Valley Sequence, whereas in Horokanai the rocks and the Yezo Group. In the southern area Cretaceous Yezo Group which is characterized by develop several shear zones of a few centimeter intercalated accumulation of submarine clastics width disturbing the ophiolite stratigraphy (Fig. 2) may be regarded as the forearc-basin sediments as well as structure (Fig. 5). accompanied by the Kamuikotan subduction. As will be described in later chapters, 'the Therefore, the tectonic emplacement of the Horo- Horokanai ophiolite contains no evidence of the kanai ophiolite onto the Kamuikotan complex as a high-pressure metamorphism, but exhibits the nappe may be dated to be in a late Mesozoic typical low-pressure metamorphism. It follows that time. the association of the Horokanai ophiolite (the low-pressure type) and the Kamuikotan metamor- phic rocks (the high-pressure type) in the study 3.PETROGRAPHY AND MINERAL ZONES area, which cannot have been formed side by side 3-1. Mafic Rocks in situ, has to be ascribed to the tectonic process. The mafic rocks of the Horokanai ophiolite dis- The structural relations of these two types of play significant changes in mineral paragenesis metamorphic rocks are shown in the cross-section (Fig. . 7) in response to changing metamorphic of Fig. 1C. This structure is not an anticlinorium grade, and the following mineral zones are disting- that is the view held in common by previous in- uished in the ascending order of metamorphic vestigators, but a nappe, possibly modified by re- grade; gional folding with a south-north axial trend Zone A: zeolites-chlorite-albite-pumpellyite plunging north. The tectonic emplacement of the Zone B: albite-chlorite-actinolite-epidote ophiolite onto the Kamuikotan metamorphic rocks Zone C: albite-oligoclase-actinolite-hornblende- 10 HIDEOISHIZUKA Uryu- Kamuikotan River Metamorphic Rocks Neogene /olcani <---bedding plane ---schistosity Horokanai ---banding plane ---layering plane Fig. 5. Structural map of the Horokanai ophiolite (after Ishizuka, 1980a). Igneous and metamorphic petrology of the Horokanai ophiolite 11 3B). The boundaries between the two zones are WEST EAST gradational, and their thickness are highly depen- Ophiolite dent upon the size of the pillow body. UK.M.R' (A) The igneous minerals are, however, partially to pervasively replaced by metamorphic minerals; ? i.e. plagioclase phenocrysts and laths are replaced by zeolites or albite that are associated with minor pumpellyite or calcite, olivine phenocrysts by chlo- rite with minor calcite and/or pumpellyite, ophitic (B) LK.M.R to subophitic clinopyroxene by chlorite, and Cr- spinel included in pseudomorphs after olivine or plagioclase phenocrysts by Cr-rich chlorite. Inter- stitial glass is altered to chlorite with disseminated fine-grained sphene and Fe-Ti oxide dust. Frac- tures and veins are wholly filled by zeolites, chlo- (C) K.M.R rite, albite, pumpellyite, calcite and rarely quartz. Mixed-Up Zone Basaltic interpillow matrix is extensively recrystal- lized to the metamorphic minerals given above, K.M.R.: Kamuikotan Metamorphic Rocks and this matrix sometimes contains chloritized or palagonitized glass shards (Plate 3C). Fig. 6. A model for the emplacement of the Horo- Zeolites are stilbite, chabazite, laumontite, kanai ophiolite (after Ishizuka, 1980a). A: wairakite, natrolite, analcime and thomsonite, as Ophiolite emplacement onto the Kamuiko- determined optically and by X-ray diffraction. tan metamorphic rocks, B: Formation of This zone can be further divided into three sub- thrust between the ophiolite and the e s s Kamuikotan metamorphic rocks, and divi- ance of chabazite, laumontite and wairakite (Fig. sion of the ophiolite into two slices of (1) 7); each subzone is mapped along a traverse route and (2), C: Emplacement of the slice (2) onto the earlier emplaced slice (1), catching (Fig. 8C); Chabazite subzone the various rock-types of the slice (1) in the basement of the slice (2), and formation of (A-C1) chlorite-chabazite-analcime-thomsonite "Mixed-Up Zone" at the boundary between (A-C2) chlorite-chabazite-analcime-stilbite the two slices of (1) and (2). Laumontite subzone (A-L1) chlorite-laumontite-analcime-thomsonite (A-L2) chlorite-laumontite-thomsonite-albite chlorite-epidote (A-L3) chlorite-laumontite-pumpellyite-albite Zone D: calcic plagioclase-hornblende- (A-L4) chlorite-laumontite-albite-quartz clinopyroxene-orthopyroxene (A-L5) chlorite-analcime-natrolite-thomsonite These four mineral zones are areally mapped (Fig. (A-L6) chlorite-analcime-pumpellyite 8A); an example of the accuracy of locating zone Wairakite subzone boundaries may be inspected in Fig. 8B. Most im- (A-W1) chlorite-wairakite-analcime-thomsonite portantly, the grade of metamorphism increases (A-W2) chlorite-wairakite-thomsonite-albite down the ophiolite stratigraphy, which is valid in (A-W3) chlorite-wairakite-pumpellyite-albite all the mapped area except for the southern area (A-W4) chlorite-wairakite-albite-quartz where several shear zones as described in the pre- Sphene and Fe-Ti oxide, and sometimes calcite, vious chapter develop disturbing the distribution occur as the accessories. Some of the chlorite pattern of mineral zones. minerals in the assemblages (A-C1), (A-C2) and (A-L1) are mixed-layer smectite/chlorite clay, as 3-1-1.ZoneA determined by X-ray diffraction. The upper sequence of the basaltic rocks (mainly All listed zeolite associations may be depicted pillow lava) belongs to this zone. The pillow lava graphically in the system CaAlSi2Og-Na2AlSi2Og- retains igneous texture, and there are two textural SiO2, as shown in Fig. 9. The zeolite paragenesis zones in a single pillow; a vitrophyric or variolitic of Fig. 9 is critical to the zeolite facies (e.g. rim (Plate 3A) and a holocrystalline core (Plate Miyashiro & Shido, 1970). The pumpellyite- 12 HIDEO ISHIZUKA LITHOLOGY BASALTIC ROCKS GABBROIC ROCKS metamorphic Pillow Tuff, Hyaloclastite Banded Massive minerals Lava &MassiveLava Flow Amphibolite Amphibolite chabazite laumontite wairakite stilbite analcime natrolite thomsonite pumpellyite prehnite chlorite albite epidote Ca-plagioclase actinolite hornblende clinopyroxene orthopyroxene quartz calcite sphene apatite Fe-Ti oxides ZONE B C D CS*:CHABAZITE SUBZONE;LS#:LAUMONTITE SUBZONE; WS#:WAIRAKITE SUBZONE Fig. 7. Mineral paragenesis for the Horokanai ophiolite (after Ishizuka, 1985). bearing assemblages also occur in other zeolite Ernst, 1979) and may belong to the prehnite- Figure 1. Sampling sites of hot spring and groundwater samples in central Tottori region (Southwest Japan) actinolite facies of Liou et al. (1985). Both 1977; Liou, 1979; Evarts & Schiffman, 1983). region20 and high-quality mineral water is available from the altered granite layers. All samples were collected at Since chabazite, laumontite and wairakite interpillow matrix; the nearby pillow core retains have the same composition except for H2O relict clinopyroxene and contains the chlorite- strike-slip fault with a compression axis in a WNW-ESE direction. After the M6.6 event, aftershocks were contin- wairakite-pumpellyite assemblage critical to the laumontite, n=2 for wairakite), the change from wairakite subzone. the chabazite to wairakite subzone,through the laumontite subzone, represents a progressive de- 3-1-2.Zone B hydration sequence of Ca-Al zeolites with rising -8.17%o on November 29th 2016 - except for irregular variation of the data a few months before the M6.6 event temperatures. The dehydration of laumontite to basaltic rocks (mainly hyaloclastite) and is char- wairakite + 2HzO was confirmed experimentally decrease gradually from -7.96%o on September 1st 2015 to -8.33%o on June 21st 2016. They then increase to by Liou (1971a). On the other hand, quartz occur- assemblage. There is a difference in texture be- ring in the laumontite and wairakite subzones tween the hyaloclastite matrix and enclosed pillow coexists with albite but not analcime. This sug- gests that these subzones are placed within the lized to yield nematoblastic texture defined by albite + quartz field, the higher-temperature field actinolite. The pillow fragment commonly retains than the dehydration of analcime + quartz (Liou, the Hakusan Meisui (HKM) site, located approximately 5km west of the Tottori earthquake epicenter (Fig. 1). 1971b; Thompson, 1971). with trace epidote, and clinopyroxene is extensive- In the highest-grade part of zone A occur the ly uralitized to aggregates of chlorite and actino- zeolite-free assemblages, prehnite-pumpellyite- lite. Furthermore, the modal proportion of meta- 15 °C is filtered and sealed in polyethylene terephthalate bottles and distributed on the market. Twenty-seven magnitude and have longer recurrence intervals, they cause serious damage to communities. The 2016 Tottori mer assemblage is critical to the prehnite- original rocks; actinolite tends to occur more pumpellyite facies, and the latter one has been re- abundantly in the matrix than in the pillow frag- out any preprocessing. Observed hydrogen and oxygen isotopic ratios were calibrated against our in-house water ment. Both the matrix and pillow fragment con- The 2016 Tottori earthquake. Earthquakes in subduction zones such as Japan can be classified into two water from September 2015 to July 2017. The 818O values with a typical error of 0.05-0.12%o (at 2o in STable 1) 1976) and from the East Taiwan ophiolite (Liou & ide, and rarely calcite. IgneousandmetamorphicpetrologyoftheHorokanaiophiolite 13 D ZONE Opx-free B :ZONE D ZONE C B ★:ZONE Opx-free :ZONE ： ZONE :ZONE Group HOROKANAI :Chabazite Subzone O:Laumontite Subzone 0:Wairakite Subzone Fig. 8. A: Mineral zone map of the Horokanai mafic rocks, B: Distribution of rocks with critical mineral assemblages of each mineral zone, and C: Distribution of rocks containing critical mineral assemb- lages of each subzone of Zone A (after Ishizuka,1985). 14 HIDEOISHIZUKA as discrete acicular to prismatic crystals, being less than 0.5 mm in length, which define a lineation (Plate 3H). Uncommonly, actinolite lamellae of stilbit albite about 20 μm in width occur in a host hornblende. chabazite analcime The textural relationships characterized by dis- wairakite crete grains or crystals may indicate equilibrium natrolite coexistence of albite with oligoclase, and actinolite with hornblende, respectively. It is, however, un- certain whether the core-rim relationship of albite and oligoclase and the lamella intergrowth of actinolite and hornblende represent equilibrium or disequilibrium textures. Na2Al2Si208 The optical discrimination of two amphiboles and sometimes two plagioclases is very difficult, -p am suz s u snqsse ar pue Fig. 9. Zeolite parageneses of Zone A on an anhyd- termined by using the electron-microprobe; they rous projection CaAlSizOs-Na2Al2Si2Og- are listed below: SiO2. (C-1) actinolite-albite-oligoclase-chlorite-epidote- quartz The ubiquitous distribution of actinolite, (C-2) actinolite-albite-oligoclase-chlorite chlorite, epidote and albite indicates that Zone B (C-3) actinolite-oligoclase-chlorite belongs to the greenschist facies. In the lowest- (C-4) actinolite-hornblende-albite-chlorite-epidote grade part of this zone, the epidote-free assemb- (C-5) actinolite-hornblende-albite-oligoclase- lage pumpellyite-actinolite-chlorite-albite occurs, chlorite-epidote-quartz but it is too restricted to define an areal mineral (C-6) actinolite-hornblende-albite-chlorite zone. (C-7) actinolite-hornblende-albite-oligoclase- chlorite-quartz 3-1-3. Zone C (C-8) hornblende-albite-oligoclase-chlorite This zone occurs occupying the middle to lower (C-9) hornblende-albite-oligoclase sequence of the basaltic rocks (mainly the lower- Sphene and Fe-Ti oxide are the common acces- most part of hyaloclastite, and the uppermost part sory minerals, but calcite is confined to the of banded amphibolite). The hyaloclastite exhibits assemblages (C-4) and (C-6). One of the most nematoblastic texture defined by actinolite and/or characteristic features of these assemblages is the hornblende, but the shape of pillow fragment is coexistence of albite and oligoclase, and hence we commonly transitional into the nematoblastic mat- may call this zone the peristerite zone. rix (Plate 3E). The banded amphibolite shows two With respect to the mineral facies, the distinct bands of less than 0.5 mm width; one rich assemblage actinolite-oligoclase-chlorite (C-3) is in fine-grained plagioclase (albite and/or oligoc- critical to the actinolite-calcic plagioclase facies of lase) and the other in nematoblastic hornblende Miyashiro (1961), which has been reported from (Plate 3F). Other constituent minerals include the transitional zone between the greenschist and chlorite, epidote, sphene, Fe-Ti oxide, and rarely amphibolite facies in contact metamorphic au- quartz and calcite. reoles (e.g. Abukuma Plateau: Shido, 1958; Kita- In most of the Zone C rocks, albite and oli- kami Mountains: Seki，1961;S Sierra Nevada: goclase coexist as discrete grains of less than 0.3 Loomis, 1966; Yap Islands: Shiraki, 1971; Van- mm diameter, which often gather to form pool- couver Island: Kuniyoshi & Liou, 1976) as well as shaped aggregates of less than 2 mm diameter from Mid-Atlantic Ridge (Miyashiro et al., 1971) (Plate 3G). The grain boundary between albite and from the East Taiwan ophiolite (Liou & and oligoclase is commonly sharp, and straight or Ernst, 1979). However, the majority of the Zone gently curved; a Becke line is sometimes distinct C rocks contain the peristerite pair with or with- at the grain boundary. Occasionally, oligoclase out the actinolite-hornblende pair, and in a strict rims of less than 50 μm width develop around a sense they do not belong to the actinolite-calcic few of albite grains. Also, in the medium-grade plagioclase facies. In this connection, it should be part of Zone C, actinolite and hornblende coexist noted that the use of the electron-probe microana- Igneous and metamorphic petrology of the Horokanai ophiolite 15 lyser has recently revealed the common occurr- most of the Zone D rocks belong to the amphibo- ence of the peristerite and/or actinolite- lite facies, but the highest-grade rocks with hornblende pairs in the transitional zone of con- pleochroic orthopyroxene may be placed in the tact metamorphic aureoles (e.g. Abukuma (hornblende-) granulite facies (e.g. Howie, 1965). Plateau: Tagiri, 1973, 1977; Sierra Nevada: Hietanen, 1974; Yap Islands: Maruyama et al., 3-2. Ultramafic Rocks 1982, 1983). Such finding requires a re- The cumulate rocks, occurring near the boundary examination of the mineral assemblages to define between the mafic and ultramafic rocks, are more the actinolite-calcic plagioclase facies. Therefore, or less affected by recrystallization, but locally in this thesis, the Horokanai Zone C is referred to they retain primary cumulate texture defined by simply as the transitional facies from the greensch- cumulus olivine, clinopyroxene and plagioclase ist to amphibolite facies. (Plate 4E). These cumulus phases are commonly unzoned, indicating that the Horokanai cumulate 3-1-4.ZoneD rocks belong largely to adcumulate in the sense of Zone D includes the lowermost sequence of the Wager et al. (1960). The effect of recrystallization basaltic rocks (banded amphibolite) and all the appears at places between cumulus phases as gabbroic rocks s (massive amphibolite). .Both forming fine-grained clinopyroxene and horn- banded and massive amphibolites display granob- blende with or without orthopyroxene. On the lastic to gneissose texture (Plates 4A, 4B, and other hand, orthopyroxene except for recrystal- 4C), but trace igneous clinopyroxene is present in lized fine-grained one is very rare but present in the massive amphibolite (metagabbro) as porphy- dunite and wehrlite, and it has very different crys- roclastic grain with margins and/or local patches tal habit from other cumulus phases, that is, it is recrystallized to fine-grained clinopyroxene and/or unusually coarse-grained and porphyroclastic hornblende (Plate 4D). (Plate 4F). Chemically, as described in the next The constituent minerals include calcic plag- chapter, such orthopyroxene has also unusually ioclase, hornblende, clinopyroxene, orthopy- magnesian composition in comparison with coex- roxene, sphene, Fe-Ti oxide and rarely apatite, to isting cumulus olivine and clinopyroxene, and which chlorite is added in the lowermost-grade rather resembles the orthopyroxene composition part of this zone. Calcic plagioclase is usually of the basal ultramafic rocks. These are reminis- saussuritized to aggregate of fine-grained albite + cent of relict phase derived from mantle peridotite chlorite + epidote + prehnite. The Z-axial color but not cumulus phase crystallized from magma. of hornblende changes from blue-green to brown The less serpentinized basal ultramafic rocks with increasing metamorphic grade. Metamorphic are composed of olivine and pyroxene with minor clinopyroxene first appears in the medium-grade spinel; the mineral proportion is shown in Fig. 10. part of Zone D, and its modal proportion tends to They have curvilinear grain boundaries between increase gradually with increasing metamorphic olivine crystals or between olivine and orthopy- grade. In the highest-grade part of this zone, roxene crystals (Plate 4G). The grain size of oli- metamorphic orthopyroxene develops with conspi- vine and orthopyroxene is typically coarse, while cuous pleochroism varying from salmon pink to spinel and clinopyroxene, of which the latter is pale green. confined to the upper horizon of the ultramafic The observed mineral assemblages in this rocks, are rather smaller in grain size. This kind zone are: of peridotite texture is characteristic of the protog- (D-1) calcic plagioclase-hornblende-chlorite ranular type, one of the representative textures of (D-2) calcic plagioclase-hornblende the mantle peridotite (Mercier & Nicolas, 1975). (D-3) calcic plagioclase-hornblende-clinopyroxene However, there is little sign of polygonization or (D-4) calcic plagioclase-hornblende-clinopyroxene- recrystallization of large crystals into aggregates of orthopyroxene finer grains with preferred orientation, even Fe-Ti oxide is common, but sphene is confined to though coarse-grained olivine and orthopyroxene the lower- to medium-grade part of this zone. Un- sometimes exhibit wavy extinction and kink- commonly, apatite occurs in the assemblage (D- bands. Such a textural relationship indicates that 3). the Horokanai ultramafic rocks have not under- The widespread occurrence of the calcic plag- gone intense deformation or plastic flow in the ioclase + hornblende assemblage indicates that mantle (Mercier & Nicolas, 1975). 16 HIDEO ISHIZUKA temperature stage through element re-distribution (re-equilibration) at temperatures corresponding to the granulite facies. 4. BULK ROCK AND MINERAL CHEMISTRIES 4-1. Analytical Procedure The major element analyses of bulk rocks were done following Fukuyama & Sakuyama (1976) by means of the Hitachi Model XMA-5A electron- probe microanalyser of Kanazawa University. The Bence-Albee (1968) method with correction fac- tors of Albee & Ray (1970) was used for data re- --dunite duction. The minor elements analysed are as fol- .--harzburgite lows: (1) V, Rb, Sr, Y and Zr by the Rigaku Model XRF-KG-4 X-ray fluorescence analyser of orthopyroxenite Kanazawa University, using the method of Nor- rish & Chappell (1967), Ikeda & Banno (1972) and Nakagawa (1978, personal communication), >Cpx 10 and (2) Cr and Ni by the Hitachi Model AAS-170- 30 atomic absorption spectrophotometry of Hok- kaido University, based on the calibration curves Fig. 10. Mineral proportion of the Horokanai ultra- constructed by artificial standard solutions. The mafic rocks (after Ishizuka, 1980a). Abbre- standard samples JB-1, JG-1 and BHVO-1 were viations: Ol=olivine, Opx=orthopyroxene, Cpx=clinopyroxene. used to check the accuracy of the analyses; both newly analysed and recommended values of JB-1 quoted for comparison are listed in Table 1. The main features of serpentinization include: Chemical compositions of selected minerals (1) olivine splits into fine, rounded grains floating were also determined using the same electron- in a network of serpentine and magnetite, (2) probe microanalyser as that of the major element orthopyroxene is almost entirely altered to bastite, analyses of bulk rocks. Supplementary micro- (3) spinel with translucent brown color is little probe analyses were done using the JEOL Model altered, but its margin is sometimes armoured by JAX-50A of Hokkaido University, and the JEOL Cr-rich chlorite, and (4) clinopyroxene is relatively Model JAX-5A of Kochi University; both micro- fresh with a trace of chlorite rinds. In addition to probes using the same data correction procedure these altered minerals, talc and calcite occur also. as that of the Hitachi micro-probe. Furthermore, sodatremolite develops in the dunite layers of 1-3 m thick within the massive harzbur- 4-2. Bulk Rock Chemistry gite. It is usually associated with spinel grains, and As many of major and minor elements are mobile Occurs as a stout prismatic, euhedral to subhedral, during such secondary processes as weathering, crystal less than 3 mm in length (Plate 4H). The hydrothermal activity, and low-grade metamorph- optical properties of the sodatremolite are as fol- ism of the zeolite to greenschist facies (Hart, lows; colorless in thin section, 2Vx(mean)=80%, 1970, 1973; Thompson, 1973a; Hajash, 1975; Bis- r<v weak, an angle between c and Z=15°, and choff & Dickson,1975; Scott & Hajash，1976), b=Y. Serpentine minerals are chrysotile and lizar- the primary compositions of the1 Horokanai dite, and no antigorite is detected in the analysed ophiolite, especially its upper basaltic rocks that samples except one sample that contains antigorite evidently suffered the low-grade metamorphism, cross-cut by a veinlet of chrysotile or lizardite. may have been masked. However, recent inves- Olivine, pyroxene and spinel described above tigations on the chemistry of volcanic rocks have are primary minerals once equilibrated with some emphasized that the elements of Ti, V, Cr, Ni, Y, high-temperature magma, but thei chemical com- Zr and the ratio of FeO*/MgO (FeO* means total positions, as will be described in the next chapter, iron as FeO) are resistant to secondary processes, have been substantially modified from the high- and their abundances can be used to estimate the Igneous andmetamorphicpetrology of theHorokanaiophiolite 17 Table 1. Analyses of chemical standard JB-1, run during analyses of samples from the Horokanai ophio- Horokanai Ophiolite lite, compared to the recommended values of 0: basaltic rocks Ando et al. (1974). : gabbroic rocks Present Value Recommended Ocean-Floor M.V** (S.D**) Value (): basaltic rocks (wt.%) /: gabbroic rocks SiO2 53.86 (0.20) 53.51 TiO2 1.34 (0.03) 1.37 Al2O3 14.86 (0.19) 14.90 % FeO* 8.05 (0.15) 8.31 MnO 0.16 (0.02) 0.15 W MgO 7.87 (0.13) 7.97 CaO 9.37 (0.08) 9.47 Na2O 2.88 02 (0.04) 2.87 K2O 1.41 (0.05) 1.48 Total 99.80 100.00 (ppm) V 214 (7) 211 Cr 409 (7) 405 Ni 139 (2) 135 Rb 39 (2) 41 Sr 440 (5) 435 人 24 (2) 26 Zr 150 (3) 153 0.5 2.0 FeO* means total iron as FeO. M.V** and (S.D**) represent mean value and standard FeO* deviation of five duplicate analyses, respectively. Fig. 11. Variation of TiO2 content against FeO*/ MgO ratio for the mafic rocks (after Ishizu- ka, 1981a). Data for ocean-floor basaltic and gabbroic rocks are from Muir & Tilley primary igneous chemistry (Cann, 1970; Pearce & (1964), Engel et al. (1965), Melson et al. Cann, 1973; Pearce, 1975; Herrmann et al., 1974; (1968), Miyashiro et al. (1969), Kay et al. Humphries & Thompson, 1977). In the following, (1970), Shido et al. (1971), Thompson et al. (1972), Campsie et al. (1973), Thompson such elements and ratio are used to characterize (1973b), Frey et al. (1974), Engel & Fisher the geochemical nature of the Horokanai (1975), Blanchard et al. (1976), and Lang- ophiolite. muir et al. (1977). A total of 30 samples (11-pillow lavas, 4- banded amphibolites, 5-massive amphibolites, 5- dolerite dikes, 2-pillow fragments in hyaloclastites, versus FeO*/MgO in Fig. 12. In those plots, the and 3-plagiogranite dikes) from the Horokanai FeO*/MgO value as the horizontal axis represents ophiolite were analysed for major elements, an indicator of the advance of fractional crystal- among which 21 samples were further analysed for lization (Shido et al., 1971). minor elements. The results of major and minor The basaltic rocks have the following sys- element analyses are listed in Tables 2 and 3, re- tematic relations; the successive increases in spectively. FeO*/MgO are accompanied by the successive in- A plot of TiO2 versus FeO*/MgO for the creases in TiO2 and V, and by the successive de- basaltic rocks (pillow lavas, banded amphibolites, creases in Cr and Ni. These compositional rela- dolerite dikes, and pillow fragments in hyaloclas- tions are the trends controlled by fractional crys- tites) and gabbroic rocks (massive amphibolites) is tallization. The relatively low FeO*/MgO values, illustrated in Fig. 11, and those of V, Cr and Ni high Cr and Ni contents indicate early-stage frac- former has lower Cr and Ni contents, higher TiO2 exist between the basaltic and gabbroic rocks; the same data as Fig. 11. Small but significant differences in chemistry basaltic and gabbroic rocks are based on the basalts of oceanic islands (Engel & Engel, 1970). Ishizuka, 1981a). Fields for 0cean-floor sented here is very different from that in alkali FeO*/MgO ratio for the mafic rocks (after It is also clear that the differentiation trend pre- Fig. 12. Variations of V, Cr and Ni contents against Miyashiro (1974) and Miyashiro & Shido (1975). tholeitic but not calc-alkaline in the sense of suggest that the fractionation trend is essentially Fe0*/Mg0 sh s m A u sn 2.0 1.5 1.0 0.5 nocrysts. On the other hand, the systematic in- as inclusion in the olivine and/or plagioclase phe- fluenced by the crystallization of spinel that occurs 100 uncertain to what extent Cr concentration are in- rophenocrysts in the pillow lavas. It is, however, N spectively, both occurring as phenocrysts or mic- 200 magnesian minerals are olivine and pyroxene, re- lavas, the large fractionation effects for Ni and Cr 000 tionation (Thompson et al., 1972). In the pillow 400 Ishizuka, 1981a). FeO*/MgO ratio for the mafic rocks (after Fig. 13. Variations of Y and Zr contents against Fe0*/Mg0 200 2.0 1.5 1.0 0.5 3 400 (ppm 20 600 Y (ppm) 0 800 : gabbroic rocks : basaltic rocks Ocean-Floor : gabbroic rocks ●: basaltic rocks ●: basaltic rocks 20 Horokanai Ophiolite 100 (pp 09 m 300 80 007 Zr (ppm) 100 500 HIDEOISHIZUKA 18 amphibolites, 16-20 of massive amphibolites, 21-25 of dolerite dikes, 26-27 of pillow fragments in hyaloclastite, and 28-30 of plagiogranite dikes. SampleNo. 01 02 03 04 05 06 07 08 09 10 11 12 13 14 15 SiO2 49.36 49.53 49.82 49.88 49.98 50.13 50.64 51.06 51.57 52.30 52.77 48.51 48.95 49.56 49.98 TiO2 1.44 1.05 1.39 1.25 1.25 1.23 1.14 1.27 1.47 1.11 1.12 1.23 1.22 1.33 1.28 Al2O3 13.47 13.29 14.98 12.61 14.25 14.14 15.45 15.44 14.44 14.48 13.57 17.22 13.87 15.92 14.71 FeO* 10.26 10.77 10.20 10.69 12.38 10.90 10.23 9.40 12.23 9.18 10.69 10.30 11.45 10.46 10.32 0.18 0.22 0.13 0.17 0.18 0.23 0.22 0.24 0.15 0.21 0.19 0.18 0.24 0.21 ous MnO 0.22 and MgO 8.51 9.65 8.42 9.75 9.53 8.72 7.92 7.48 7.41 9.28 9.17 7.95 11.54 8.55 8.48 12.42 7.81 9.88 11.07 10.27 7.90 9.87 8.48 mei CaO 13.24 12.66 11.71 12.19 10.12 9.96 12.74 Na2O 3.49 3.46 4.23 3.44 4.15 3.65 3.27 3.92 4.39 3.96 4.39 2.06 2.43 3.18 2.27 K2O 0.14 0.05 0.01 0.04 0.02 0.43 0.04 0.72 0.28 0.32 0.42 0.28 0.12 0.64 0.07 Total 100.09 100.68100.89 100.25 99.55 99.30 99.99 99.78 99.93 100.65 100.83 99.93 99.88 99.84 100.06 peti FeO*/MgO 1.21 1.12 1.21 1.10 1.30 1.25 1.29 1.26 1.65 0.99 1.17 1.30 0.99 1.22 1.22 rology SampleNo. 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 100 SiO2 47.18 47.71 48.18 49.54 49.66 49.53 50.82 51.13 51.89 52.44 50.77 50.97 66.29 67.03 69.04 Ho TiO2 0.63 0.83 0.50 0.73 0.67 1.33 1.28 1.26 1.17 1.23 1.09 0.99 0.89 0.71 0.63 Al2O3 14.64. 13.80 14.08 14.29 14.48 13.65 15.21 14.41 14.49 13.66 14.93 15.21 14.79 14.80 13.05 FeO* 8.47 10.45 8.94 9.48 8.86 12.35 9.43 9.99 9.69 11.61 10.16 9.52 6.97 6.05 4.97 MnO 0.28 0.24 0.21 0.20 0.15 0.22 0.22 0.19 0.20 0.28 0.19 0.20 0.03 0.02 0.01 MgO 13.51 10.86 13.79 9.53 10.41 8.15 7.64 7.43 7.48 7.60 9.95 9.68 1.27 1.01 0.75 CaO 12.58 12.82 12.10 13.12 12.86 10.73 9.86 11.47 10.76 8.92 8.10 8.08 3.79 4.01 4.44 Na2O 1.66 2.16 1.45 2.05 1.98 3.21 3.94 3.19 4.08 3.98 3.73 4.67 5.89 5.70 6.78 K2O 0.22 0.28 0.10 0.13 0.17 0.28 0.63 0.57 0.79 0.08 0.27 0.14 0.05 0.10 0.04 Total 99.17 99.15 99.35 99.07 99.24 99.45 99.03 99.64 100.55 99.80. 99.19 99.46 99.97 99.43 99.71 FeO*/MgO 0.63 0.96 0.65 0.99 0.85 1.52 1.23 1.34 1.30 1.53 1.02 0.98 5.49 5.99 6.63 FeO*meanstotal iron asFeO. 6 20 HIDEOISHIZUKA Table 3. Minor element contents of representative rock-types from the Horokanai ophiolite (in ppm). Sample Nos. correspond to those in Table 2. Sample No. 01 03 04 05 06 08 09 11 12 13 14 V 305 318 283 335 300 324 413 325 351 270 337 Cr 339 355 289 350 329 363 289 368 374 410 385 Ni 139 115 144 110 121 149 76 125 130 161 152 7 6 Rb 2 4 1 2 6 10 1 2 Sr 65 170 269 153 278 162 747 171 141 129 103 Y 32 34 29 37 33 35 45 31 30 27 37 Zr 70 75 74 75 69 68 98 69 71 67 78 Sample No. 16 17 18 19 20 21 22 23 28 30 V 93 261 110 213 178 375 363 380 35 22 Cr 973 405 828 423 655 293 325 327 19 12 Ni 442 193 317 185 274 85 130 135 8 5 Rb 2 2 1 2 3 5 3 1 4 3 Sr 78 52 65 62 57 305 123 81 145 121 人 23 26 28 31 27 41 35 36 104 113 Zr 45 48 53 58 50 90 70 80 231 219 than the latter dose.However, the contents of rocks fall in the abyssal tholeite field. Furth- these elements gradually change from the gabbroic ermore, the abundances of Ti, Cr and Ni are sig- to basaltic rocks. The similar trends are also dis- nificantly lower in island arc and higher in oceanic tinct for Y and Zr as illustrated in Fig. 13. Y and island than in abyssal tholeites (e.g. Jakes & Zr are notable “incompatible” elements that are White, 1972). Some useful diagnostic diagrams as largely rejected by crystallizing phases and con- centrate in the residual liquid, and there is usually a positive correlation between these “incompati- ble' elements and the FeO*/MgO ratio in frac- A.D:Island Arc Tholeite tionated rock series (Erlank & Kable,1976; Tar- 2.0 B.D:Abyssal Tholeite ney et al., 1977). These facts suggest that the gab- C.D:Calc-Alkali Basalt broic rocks now presented 1by the massive ●: Horokanai Ophiolite amphibolites were co-magmatic in origin with the 1.5 basaltic rocks, the former representingless- E fractionated or more-primitive parts of the basaltic magma. The plagiogranite dikes have distinct geoche- 1.0 001/!1 mistry as compared with the basaltic and gabbroic rocks; Fe0*/MgO (5.49-6.63) value are appreci- ably high, Y (104-113 ppm) and Zr (219-231 ppm) 0.5 are relatively enriched, and Cr (12-19 ppm) and Ni (5-8 ppm) are strongly depleted. These chemi- cal natures are typical of oceanic plagiogranite, a collective term proposed by Coleman & Peterman 50 100 150 200 (1975). (wdd）z On the other hand, the Horokanai basaltic rocks are plotted in the Ti versus Zr and Ti-Zr-Y Fig. 14. The Horokanai basaltic rocks plotted in the basalt discrimination diagrams of Pearce & Cann Ti-Zr basalt discrimination diagram (after (1973) in Figs. 14 and 15, respectively. From these Ishizuka, 198la). Fields for various basalt- diagrams, it is obvious that most of the basaltic types are after Pearce & Cann (1973). Igneous andmetamorphicpetrologyof theHorokanai ophiolite 21 1 : Oceanic Island or Continental Tholeiite Ti/100 2.3: Island Arc Tholeiite 3 : Abyssal Tholeiite 3.4: Calc-Alkali Basalt ●：Horokanai Ophiolite Ti/100 Zr Y.3 50 50 Zr 50 Y.3 Fig. 15. The Horokanai basaltic rocks plotted in the Ti-Zr-Y basalt discrimination diagram (after Ishizuka, 1981a). Fields for various basalt-types are after Pearce & Cann (1973). illustrated in Fig. 16 are obtained by plotting Cr, occurrence of the last-type is commonly restricted Ni versus Ti, which indicates that the Horokanai to the variolitic zone of the pillow lavas. The mic- basaltic rocks fall in the field of the abyssal rophenocryst is unaltered (rarely altered to a trace tholeites, but not in the fields of the island arc of chlorite or actinolite around the margin and/or nor oceanic island tholeites. It follows that the along the cracks), whereas the groundmass and Horokanai ophiolite has a close affinity with abys- dendritic types are often replaced by chlorite or sal tholeiites, and that an oceanic environment for actinolite. its generation is more suitable than an island arc All pyroxene analyses are plotted in the or oceanic island setting. pyroxene quadrilateral, as shown in Fig. 17, even thought some contain appreciable amounts of Ti 4-3. Relict Mineral Chemistry and Al. The pyroxenes are mainly augites, extend- Relict mineral 1analyseswereperformedfor ing from diopsidic or salitic pyroxene to augite, pyroxene and spinel, of which representative and no orthopyroxene, pigeonite and subcalcic au- analyses are listed in Tables 4 and 5, respectively. gite occur in any of the sample analysed. The microphenocryst t pyroxenes are often 4-3-1. Pyroxene (Table 4) chemically zoned, showing, from the core to the Relict pyroxenes in the Horokanai pillow lavas are rim, Ca and Mg depletion and Fe* (total Fe as divided into three types; (1) microphenocryst and Fe2+) enrichment trends. As seen in this trend, (2) groundmass pyroxene based on grain size, and the rims are strongly depleted in Cr, relative to (3) dendritic pyroxene based on shape. The the cores, even if the relative Fe* enrichment is 22 HIDEO ISHIZUKA Cr (ppm) 1000F 500 100 50 .. : Island Arc Tholeiite (-': Abyssal Tholeite 二"!:OceanicIslandTholeite :HorokanaiOphiolite Ni (ppm) 500 100 50 0.5 1.0 1.5 Ti/10000 (wda) Fig. 16. The Horokanai basaltic rocks plotted in Cr-Ti and Ni-Ti basalt discrimination diagram (after Ishizu- ka, 1981a). Data for island arc tholeites are from Coats (1952), Nockolds & Allen (1956),Stark (1963), Taylor et al. (1969), Gill (1970), Lowder & Carmichael (1970), Evart & Bryan (1972), Evart et al. (1973) and Shiraki et al. (1977), those for abyssal tholeites from Muir & Tilley (1964), Engel et al. (1965), Melson et al. (1968), Kay et al. (1970), Thompson et al. (1972), Campsie et al. (1973), Frey et al. (1974), Engel & Fisher (1975), Blanchard et al. (1976), Langmuir et al. (1977), Flower et al. (1979), Bougault et al. (1979) and Wood et al. (1979), and those for 0ceanic island tholeites from Wager & Michell (1953), Kuno et al. (1957), Nockolds & Allen (1959), Bartel et al. (1963), Macdo- nald & Eaton (1964), McBirney & Williams (1969), Gunn (1971) and Wood (1978). Igneous andmetamorphicpetrology of the Horokanai ophiolite 23 Table 4. Representative analyses of relict pyroxenes in the Horokanai pillow lavas. Sample Nos. correspond to those in Table 2. Sample No. MicrophenocrystType 01 03 06 09 core-rim core-rim core-rim core-rim SiO2 50.13 52.07 49.98 52.36 51.01 51.67 50.53 51.98 TiO2 0.65 0.65 0.66 0.54 0.71 0.46 0.68 0.56 AlO 4.19 2.58 4.11 2.88 4.25 3.15 4.37 2.99 CrO 1.14 0.10 0.91 0.18 0.65 0.28 0.50 0.30 FeO* 6.38 10.12 6.11 8.45 6.47 8.85 7.58 8.99 MnO 0.13 0.23 0.11 0.18 0.15 0.15 0.15 0.25 MgO 15.97 14.67 15.85 15.24 15.81 15.18 15.47 U-Pb ages from this formation indicate a Late Silurian-Early Devonian CaO 21.09 19.73 21.17 20.85 21.55 20.89 20.83 20.24 NaO 0.23 0.35 0.25 0.16 0.21 0.14 0.22 0.26 Total 99.91 100.50 99.15 100.84 100.81 100.77 100.33 100.73 Atomicproportions.O=6.0 Si 1.856 1.931 1.862 1.924 1.869 1.907 1.866 1.917 Ti 0.018 0.018 0.018 0.015 0.020 0.013 0.019 0.016 A1 faulted contact with the Ichinotani Formation, and their original strat- 0.113 0.181 reported from this formation (e.g., Suzuki and Shino, 2020; Tazawa 0.184 0.137 0.190 0.130 Cr 0.033 0.003 0.027 0.005 0.019 0.008 0.015 0.009 Fe 0.198 0.314 0.190 0.260 0.198 0.273 Paleozoic strata of the Hida Gaien belt consist mainly of Ordovician 0.277 Mn 0.004 0.007 0.003 0.006 0.005 0.005 0.005 0.008 Mg 0.881 0.811 0.880 0.835 0.863 0.835 0.851 0.833 Ca 0.837 0.784 0.845 0.821 0.846 0.826 0.824 (1957) in the southern part of this area (Fig. 2B), consists of Permian Na 0.017 0.025 0.018 0.011 0.015 0.010 0.016 0.019 Sample No. Micropheno.Type GroundmassType 10 01 0 06 60 10 core-rim SiO2 50.16 52.15 51.26 50.88 49.88 50.93 50.98 TiO 0.73 0.55 0.51 0.64 0.76 0.56 0.57 AlO 4.25 2.79 3.22 2.35 4.21 2.65 2.73 CrO 0.45 0.10 0.25 0.06 0.40 0.10 0.03 NNE-SSW and dip 30°-80° to the northwest. Yoshida and Tazawa 8.33 10.24 8.56 10.74 7.59 9.45 igraphic relationship might have been conformable or exhibited a minor MnO 0.23 0.26 0.18 0.28 0.13 0.28 0.34 MgO 15.89 15.02 The Sorayama Formation conformably overlies the Mizuyagadani 15.47 14.97 15.58 14.27 CaO 20.01 19.42 20.56 19.88 21.56 20.59 19.83 NaO 0.28 0.24 0.26 0.31 0.31 0.28 0.33 Total 100.33 100.77 101.02 100.61 99.81 100.42 100.24 Atomic proportions,O=6.0 Si 1.858 1.927 1.886 1.898 1.859 1.896 1.910 Ti foraminiferas indicate that this formation is Early to Late Carboniferous 0.015 0.014 0.018 0.021 0.016 0.016 A1 0.186 0.122 0.140 clastic rocks, Devonian limestone, Carboniferous mafic and felsic vol- 0.185 0.116 0.121 Cr 0.013 0.003 0.007 0.002 0.012 0.003 0.001 Fe 0.258 0.316 0.263 0.335 0.237 0.294 0.350 Mn 0.007 0.008 0.006 0.009 0.004 0.009 0.011 Mg 0.877 0.827 0.890 0.831 0.864 0.797 Ca 0.794 0.769 0.811 0.795 0.861 0.821 0.796 Na 0.020 0.017 0.019 0.022 0.022 including calcareous concretions from which Early Silurian radiolarians 0.024 24 HIDEO ISHIZUKA (Table 4. continued) Sample No. Dendritic Type 04 05 08 SiO2 52.78 52.14 52.12 51.88 51.97 TiO2 0.67 0.58 0.49 0.52 0.56 Al2O3 2.38 2.24 2.42 2.15 2.22 Cr2O3 0.11 0.04 0.07 0.04 0.08 FeO* 10.12 8.55 8.56 9.47 10.22 MnO 0.22 0.18 0.21 0.23 0.17 MgO 15.98 17.14 17.33 17.51 16.54 CaO 18.42 19.37 18.66 18.62 18.67 Na2O 0.23 0.27 0.32 0.26 0.54 Total 100.91 100.51 100.18 100.68 100.97 Atomic proportions, O=6.0 Si 1.939 1.919 1.921 1.912 1.917 Ti 0.019 0.016 0.014 0.014 0.016 Al 0.103 0.097 0.105 0.093 0.097 Cr 0.003 0.001 0.002 0.001 0.002 Fe 0.311 0.263 0.264 0.292 0.315 Mn 0.007 0.006 0.007 0.007 0.005 Mg 0.875 0.940 0.952 0.962 0.909 Ca 0.725 0.764 0.737 0.735 0.738 Na 0.016 0.019 0.023 0.019 0.039 FeO* means total iron as FeO. ·,o : Microphenocryst Type (solid:core, open: rim ) + : Groundmass Type : Dendritic Type n 20 30 Di 40 Fig. 17. Relict pyroxene compositions in the pillow lavas plotted in the pyroxene quadrilateral (after Ishizu- ka, 1981a). Tie line represents the core and the rim in a zoned microphenocryst pyroxene. Igneous andmetamorphicpetrologyof theHorokanai ophiolite 25 1.5 1.0 (wt. Cr203 0.5 ·,o :Microphenocryst Type (solid:core , open:rim) +:Groundmass Type D :Dendritic Type 0.1 0.2 0.3 0.4 Fe*/(Fe* + Mg) Fig. 18. Variation of Cr2O3 content against Fe*/(Fe*+Mg) ratio for relict pyroxenes in the Horokanai pillow lavas (after Ishizuka, 1981a). Symboles with numbers (1-6) represent the step-scan analyses in a rep- resentative zoned microphenocryst pyroxene, of which analytical points correspond to those in a sketch illustrated in this figure. still meager. The typical tendency for Cr depletion (Fig. 6D) and a small amount of limestone. In the sandstone-mudstone is illustrated in Fig. 18, which shows that the mic- with the microphenocryst pyroxenes (some of the rophenocryst pyroxenes with Fe*/(Fe*+Mg) ratio yellowish white in color and is intercalated with black siliceous less than 0.25 exhibit rapid depletion trends in Cr, such as fusulinids and crinoid stems, which are filled with calcareous Umeda and Ezaki (1997) is included in this unit. The felsic tuff is subdivide this formation into units MB1, MB2, MB3, MB4, and MB5 in little Cr. Such behaviour of Cr is also observed by pyroxenes follows a similar pattern to that in the the step-scan analyses from the core to the rim in microphenocrysts（ (Fig. 18), even though the fusulinids (Tazawa and Matsumoto, 1998; Ueno and Tazawa, 2004), succession, the sandstone grades gradually to mudstone in each bed, and pyroxene (Fig. 18). The strong partitioning of Cr less than 0.25 are restricted to the cores of coarser into pyroxene, relative to liquid, may explain the groundmass grains. On the other hand, the com- sharp depletion trend within a single grain; once positions change continuously from the cores of closed-system crystallization ensues (i.e.subsequ- eous sandstone. These rocks strike WNW-ESE to ENE-WSW and dip ence to eruption), pyroxene crystallization rapidly ones, through the rims of the microphenocryst depleted residual liquid in Cr, and the liquid was pyroxenes (Fig. 17), again indicating the tholeiitic effectively depleted in Cr by the stage at which felsic tuff, tuffaceous sandstone and mudstone, and siliceous mudstone. Unit MB2 consists of conglomerate, sandstone, and limestone. This ma, from which these pyroxenes crystallized, was crystallized. Such a strong depletion trend in Cr of a tholeitic basalt being rich in Cr. stone is blueish gray in color and contains ooids and abundant bioclasts, On the contrary, the chemical compositions alkaline basalts (Dungan et al., 1979; Schweitzer et The Moribu Formation consists of felsic tuff, tuffaceous sandstone al., 1979). from those of the microphenocryst and ground- The groundmass pyroxenes are commonly mass ones; the former is poor in Ca and rich in 26 HIDEOISHIZUKA Mg as compared with the latter (Fig. 17). The ly included in the pseudomorphs after olivine or compositional trend of the dendritic pyroxenes di- occasionally plagioclase phenocrysts, never H6 M8 trend of pyroxenes. Furthermore, it is noteworthy euhedral to subhedral and from 5 to 30 μm across that the Cr content in the dendritic pyroxenes is in size, sometimes reaching the maximum width of H4 100 μm across. Their margins are often altered to groundmass ones at equivalent Fe*/(Fe*+Mg) magnetite or Cr-rich chlorite (Cr2Os=0.5-1.0 ratio and shows no sign of depletion trend (Fig. -M8 18). These differences, as discussed by Bryan brown to pale brown in color. (1972a), can be explained by the disequilibrium In general, spinel is very sensitive to subsoli- crystallization of the dendritic pyroxenes caused dus re-equilibration and thus its chemistry in by the supercooling. The spherulitic, skeletal and holocrystalline rocks (for example,an inner core dendritic textures are commonly observed in the of pillow body) must be interpreted with caution variolitic zone of the Horokanai pillow lavas, and due to the possibility of reaction between early- these quenched crystal morphologies confirm the formed spinel and residual liquid (i.e. Irvine, supercooling process resulting from a submarine 1965, 1967). The Horokanai relict spinels occur as eruption (Lofgren, 1971; Bryan, 1972b). accessory crystals included in olivine or plagioclase phenocrysts in quenched zones of the pillow lavas, 4-3-2. Spinel (Table 5) and hence the effect of subsolidus re-equilibration Relict spinels are often contained in quenched R1 (10) liquidus phase to crystallize prior to or simul- pillow lavas, though only a few grains were recog- taneously with 1olivineand/orplagioclasefrom nized in a thin section. These spinels are common- basaltic liquid. Horokanai Ophiolite O, o :included in olivine (solid :core , open:rim) + :included in plagioclase (->:Abyssal Tholeiite ):Oceanic Island Tholeite :IslandArcTholeite AI Fe3+ 30 30 Cr 50 A1 Fig. 19. Variation of Fe3+-Cr-Al ratio for relict spinels in the pillow lavas (after Ishizuka, 1981a). Tie line represents the core and the rim in a zoned spinel. Spinel data for abysal tholeites are from Bryan (1972a), Frey et al. (1974), Sigurdsson & Schilling (1976), Sato et al. (1979), Templeman (1979) and Graham et al. (1979), those for oceanic island tholeites from Evans & Wright (1972), and those for island arc tholeiites from Shiraki et al. (1977). Table 5. Representative analyses of relict spinels in the Horokanai pillow lavas. Sample Nos. correspond to those in Table 2. Fe3+ and Fe2+ were calculated from total The Sorayama Formation consists mainly of andesitic volcanic rocks gray in color, is medium to coarse grained, and contains local thinly t0 05 08 11 sno 01* 01* 01* pl* *10 pl* and have been replaced by chlorite or calcite with opaque minerals. The 01* *10 01* 01* core—rim this part contain thin felsic tuff and sandstone layers that grade from me TiO2 0.52 0.45 0.43 0.44 0.72 0.50 0.76 0.47 0.51 0.50 0.56 0.47 Al2O3 22.49 22.68 29.68 23.57 21.98 30.56 21.95 30.66 28.50 22.72 24.74 22.09 Cr2O3 39.34 42.88 36.87 39.68 40.58 35.63 41.58 36.76 38.72 42.47 39.84 43.47 Fe2O3 8.37 6.76 5.02 8.04 8.11 5.36 7.72 4.02 5.05 5.35 6.28 5.83 pe FeO 13.43 11.36 10.44 12.90 15.22 9.82 14.36 10.89 10.48 12.30 12.43 11.55 MgO 14.91 15.78 17.03 15.52 13.14 17.50 13.79 16.89 17.06 15.55 15.59 15.70 d Total 99.06 99.91 99.47 100.15 99.75 99.37 100.16 99.69 100.32 98.89 99.44 99.11 the mudstone varies in thickness from several millimeters to several tens of Ti 0.012 0.010 0.009 0.010 0.017 0.011 0.018 0.010 0.011 0.012 0.013 0.011 Al 0.815 0.809 1.024 0.840 0.800 1.049 0.793 1.053 0.981 0.819 0.880 0.795 Cr 0.956 1.025 0.853 0.948 0.991 0.820 1.007 0.846 0.894 1.027 0.950 1.050 op Fe3+ 0.194 0.154 0.111 0.183 0.188 0.117 0.178 0.088 0.111 0.123 0.143 0.134 Fe2+ 0.345 0.287 0.256 0.326 0.393 0.239 0.368 0.265 0.256 0.315 0.314 0.295 Mg 0.683 0.711 0.743 0.699 0.605 0.759 0.630 0.733 0.742 0.709 0.701 0.715 glomerates contain abundant limestone clasts in a tuffaceous matrix. 乙 28 HIDEO ISHIZUKA All spinel analyses are plotted in Figs. 19 and Evans & Wright, 1972) and ocean-floor (Bryan, 20 in terms of Fe3+-Cr-Al and Cr/(Cr+Al) versus 1972a; Frey et al., 1974; Sigurdsson & Schilling, Fe2+/(Fe2+ +Mg), respectively. The Horokanai re- 1976; Sato et al., 1979; Templeman, 1979; Gra- lict spinels are characterized by low Fe3+ content, ham et al., 1979). Although the spinels occur in moderate Cr/(Cr+Al)ratio and relatively low basaltic rocks with relatively primitive nature such Fe2+/(Fe2+ +Mg) ratio. There is also positive cor- as containing high Cr and Ni bulk rock contents Fe2+/(Fe2+ + Mg) relation between and Cr/ and having low FeO*/MgO bulk rock values, their (Cr+Al) ratios. Furthermore, the spinel composi- compositional variations are remarkably different tion is related to the mode of occurrence; spinels among various tectonic settings (Figs. 19 and 20). included in plagioclase are in average slightly In particular, the spinels in the abyssal tholeites more magnesian and less ferric than those in oli- have different chemical trend from those in both vine. Although most of spinels are chemically the island arc and oceanic island tholeites; the homogeneous, weak chemical zoning is sometimes spinels in the abyssal tholeiites have evidently low- detected in coarser grains; from the core to the er Fe3+ content and Cr/(Cr+Al) ratio than those rim, an increase in Cr and decrease in Al, a minor in both the island arc and oceanic island tholeiites. increase in Fe2+ and decrease in Mg, and a slight- It is clear from Figs. 19 and 20 that the composi- ly increase in Fe3+ (Figs. 19 and 20). Composi- tional variation of the Horokanai relict spinels is tional variation as well as zoning pattern of the re- very close to that in the abyssal tholeites, lict spinels are mainly due to the extensive sub- although the aluminous spinels characteristics of stitution of Al for Cr. the high magnesian primitive tholeites found on As is shown in the bulk rock chemistry, a dis- Legs 2, 3, 37 and in the FAMOUS area have not crimination of the basaltic rocks erupted in va- been recognized in the Horokanai pillow lavas. It rious tectonic settings is also obtained by the follows again that the Horokanai basaltic pillow spinel chemistry. The spinels occur as liquidus lavas have a close chemical affinity with abyssal phase in tholeitic basalts in island arc (Mariana: tholeiites. Shiraki et al., 1977), 0ceanic island (Hawaii: 4-4. Metamorphic Mineral Chemistry Analyses of metamorphic minerals were per- 1.0 formed for zeolites, pumpellyite, chlorite, epidote, plagioclase, amphibole, pyroxene, olivine and Horokanai Ophiolite ·,0 : included in olivine spinel. Representative analyses are listed in Table (solid:core,open:rim) 0.8 7 to Table 14, and mineral assemblages of analy- +: included in plagioclase sed mafic samples are given in Table 6. 0.6 4-4-1.Zeolites (Table 7) Zeolites occurring in Zone A are commonly homogeneous in composition, approaching the 0.4 ideal stoichiometry of chabazite, laumontite, wairakite, analcime and thomsonite, respectively. However, analcime occurring in the wairakite sub- 0.2 (>' : Abyssal TH :Oceanic Island TH zone as a pseudomorph after plagioclase has a : Island Arc TH CaO content of up to 4 wt.%, and shows substan- tial solid solution toward wairakite. Such a 0.4 0.6 0.8 0.2 1.0 calcium-bearing analcime is rare, but has been de- Fe2'/(Fe2+ Mg) scribed in low-grade metamorphic terrains (Seki, 1971; Surdam,1973; Evarts & Schiffman,1983). Fig. 20. Variation of Cr/(Cr+Al) ratio against Fe2+/ 4-4-2. Pumpellyite (Table 8) (Fe2+ +Mg) ratio for relict spinels in the pil- Pumpellyite occurs in the laumontite and waira- low lavas (after Ishizuka, 1981a). Tie line kite subzones of Zone A, and shows various mod- represents the core and the rim in a zoned es of occurrence, namely, filling vein and amyg- spinel. Spinel data for abyssal, oceanic is- land and island arc tholeiites are based on dule, occupying interstitial matrix, and replacing the same data as Fig. 19. igneous minerals (Plate 5). The best crystallized Igneous andmetamorphicpetrology of theHorokanai ophiolite 29 Table 6. Mineral assemblages of analysed mafic rocks from the Horokanai ophiolite. Abbreviations; Ze=zeolite, Pu=pumpellyite, Ch=chlorite, Ab=albite, Ep=epidote, Pl=Ca-plagioclase, At=actinolite, Hb=hornblende, Cx=clinopyroxene, Ox=orthopyroxene, Qz=quartz, Cc=calcite, Sp=sphene, FT=Fe-Ti oxide, F*/M= se=- 'sd=+ ( se u lo=o) e w pb oi yo yiq oo SampleNo. Zone Ze Pu 1ChAbEp / L dS zOxOx qH 1I ( AC151 A + 一 一 + + + 1.12 AL351 A + × + 一 一 + + 1.21 AL352 A + + + + + × AL331 A + + 十 + 一 + + 1.30 AL332 A 十 + + + + + AL651 + 十 + + + AL652 A + + 十 + ALW02 A + ： + + + + AW151 A + + + AW351 A + + + + + + 1.29 AW352 A + + + AW331 A B1501 B + + + + 1.02 B1502 B + + + + + + B1503 B + + 0.98 C1501 C + 十 十 + + C2501 C + + 1.17 × × C4301 C + + + + 1.22 十 C5501 C + + + + 1.22 C6501 C + + + + 1.30 × × C7301 C + + + + + C8501 C + + + + C9501 C + + + + 0.99 x C9301 C + + + D1501 D + + 0.99 D2501 D + + + + D2502 D + + + × D2503 D + 0.96 × × D3501 D + + + D3502** D + + + 一 D3301 D + + 0.85 + D4301 D + + + D4302 D + + + 0.65 D4401 D + + 0.63 D3502** contains small amounts of apatite. 30 HIDEOISHIZUKA Table 7. Representative analyses of zeolite species from Horokanai Zone A. Sample No. Cha Lau Wai Anal Ca-Anal Thom AC151 AL351 AW351 AL651 AW151 AC151 SiO2 47.13 50.81 54.85 54.21 54.87 37.21 Al2O3 18.89 21.79 23.02 22.87 22.96 32.10 Fe203* 0.08 0.10 0.13 0.07 0.05 0.05 CaO 10.85 11.85 12.12 0.09 3.96 14.02 Na2O 0.41 0.32 0.02 13.75 9.82 3.97 K20 0.05 0.03 0.02 0.03 0.02 0.02 Total 77.41 84.90 90.16 91.02 91.68 87.37 Atomic proportions** Si 4.045 3.977 4.024 2.003 2.005 4.957 Al 1.911 2.011 1.991 0.996 0.989 5.042 Fe 0.005 0.006 0.007 0.002 0.001 0.005 Ca 0.998 0.994 0.953 0.004 0.155 2.001 Na 0.068 0.049 0.003 0.985 0.696 1.026 K 0.005 0.003 0.002 0.001 0.001 0.003 Abbreviations: Cha=chabazite, Lau=laumontite, Wai=wairakite, Anal=analcime, Ca-Anal=Ca-analcime, Thom=thomsonite. Fe2O3* means total iron as Fe2O3. Atomic proportions** were calculated on the basis of O=12 for chabazite, laumonite, and wairakite, O=6 for analcime, and O=20 for thomsonite. Table 8. Representative analyses of Zone A pumpellyites filling vein (V) and amygdule (A), and replacing plagiocla- se (P), olivine (O) and clinopyroxene (C), and occupying matrix (M). Atomic proportions were calculated on the basis of total cations=16.0. Sample No. Laumontite Subzone AL351 AL352 V A P 0 C M A P M SiO2 35.96 35.47 36.47 35.25 35.80 34.96 35.53 37.77 35.59 Al2O3 11.40 12.43 16.44 13.35 14.64 5.13 12.13 18.78 12.07 FeO* 23.50 24.61 16.38 22.47 20.21 32.01 21.20 11.72 22.05 MnO 0.14 0.15 0.11 0.29 0.57 0.14 0.09 0.19 0.15 MgO 1.96 0.80 3.06 1.65 2.15 1.44 2.55 3.45 2.07 CaO 21.91 21.37 22.40 21.73 22.05 21.39 21.92 22.57 21.86 Total 94.87 94.83 94.86 94.74 95.42 95.07 93.42 94.48 93.79 Atomic proportions !S 6.02 5.98 5.94 5.89 5.89 6.02 5.99 6.09 6.00 Al 2.25 2.47 3.16 2.63 2.84 1.04 2.41 3.57 2.40 Fe 3.29 3.47 2.23 3.14 2.78 4.61 2.99 1.58 3.11 Mn 0.02 0.02 0.02 0.04 0.08 0.02 0.01 0.03 0.02 Mg 0.49 0.20 0.74 0.41 0.53 0.37 0.64 0.83 0.52 Ca 3.93 3.86 3.91 3.89 3.88 3.94 3.96 3.90 3.95 Igneous and metamorphic petrology of the Horokanai ophiolite 31 (Table 8. continued) Sample No. Laumonite Subzone For unit MB3, the compositions plot in a broad range from the oceanic AL331 V P 0 M V P 0 C M SiO2 35.42 35.53 35.77 35.54 35.38 37.16 36.08 36.17 34.85 Al2O3 9.53 19.04 15.50 11.75 10.57 18.30 13.42 16.61 6.49 FeO* 25.95 14.57 18.84 25.01 24.17 13.60 axes and contents of the measured elements as the vertical axes are 17.72 31.06 MnO 0.10 0.19 0.43 0.11 0.13 0.07 0.12 0.35 0.22 MgO 1.83 1.74 1.89 0.80 1.89 2.82 2.32 2.09 0.86 CaO 21.91 22.01 21.36 21.91 21.50 22.91 22.46 22.03 21.32 Total 94.74 93.08 93.79 95.12 93.64 94.86 95.28 94.97 94.80 Atomic proportions Si 5.99 5.89 5.97 5.98 6.02 6.01 5.95 5.93 6.01 Al 1.90 3.72 3.05 2.33 2.12 3.49 2.61 3.21 1.32 Fe 3.67 2.02 2.63 3.52 3.44 1.84 2.88 2.43 4.48 Mn 0.01 0.03 0.06 0.02 0.02 0.01 0.02 0.05 0.03 Mg 0.46 0.43 0.47 0.20 0.48 0.68 0.57 0.51 0.22 Ca 3.97 3.91 3.82 3.95 3.92 3.97 3.97 3.87 3.94 Sample No. Laumonite Subzone WairakiteSubzone AL651 AL652 AW351 V A 0 A P M V A P SiO2 35.63 34.89 37.08 36.61 the samples. The volcanic rocks of the Sorayama Formation (SR1 and 35.26 35.92 36.80 37.46 Al2O3 14.48 15.20 16.25 17.43 active continental margin fields (Fig. 11A and B). Data from unit MB2 of 10.63 14.40 18.01 20.85 FeO* 21.70 20.50 14.80 16.00 13.01 25.48 18.79 12.95 10.22 MnO 0.15 0.06 0.50 0.17 0.16 and HREEs (Gd-Lu) than observed in other units. For example, in sample 0.16 0.21 0.15 MgO 1.17 1.23 2.82 1.92 2.85 1.72 2.37 3.39 3.25 CaO 21.87 21.25 22.71 22.19 22.43 21.61 22.11 22.74 23.35 Total similar with those of the Sorayama Formation in terms of their REE 93.13 94.16 94.32 94.19 94.90 93.75 94.10 95.28 Atomic proportions !S 5.93 5.90 6.08 6.02 6.03 5.94 5.99 5.98 5.96 Al 2.84 3.03 3.14 3.38 3.58 2.11 2.82 3.45 3.91 Fe 3.02 2.90 2.03 2.20 1.77 3.59 2.62 1.76 1.36 Mn 0.02 0.01 0.07 0.02 0.02 0.03 0.02 0.03 0.02 Mg 0.29 0.31 0.69 0.47 0.69 0.43 0.59 0.82 0.77 Ca 3.90 3.85 3.99 3.91 3.91 3.90 3.95 3.96 3.98 32 HIDEOISHIZUKA (Table 8. continued) Sample No. WairakiteSubzone AW351 AW352 AW331 C M V A 0 M SiO2 37.03 37.56 36.32 35.76 36.78 37.52 37.11 35.79 35.72 Al2O3 16.49 19.14 13.32 17.77 15.62 21.34 17.26 11.57 20.66 FeO* 13.87 10.31 19.57 15.19 17.28 9.60 14.30 22.97 13.14 MnO 0.30 1.48 0.05 0.23 0.12 0.22 0.24 0.19 0.15 MgO 4.06 3.44 2.97 2.14 2.46 3.03 3.14 2.05 2.21 CaO 22.75 23.21 22.52 21.97 22.54 23.30 22.73 22.29 21.65 Total 94.50 95.14 94.75 93.06 94.80 95.01 94.78 94.86 93.53 Atomic proportions !S 6.00 6.01 5.99 5.94 6.03 5.98 6.02 5.98 5.85 Al 3.15 3.61 2.59 3.48 3.02 4.01 3.30 2.28 3.99 Fe 1.88 1.38 2.70 2.11 2.37 1.28 1.94 3.21 1.80 Mn 0.04 0.20 0.01 0.03 0.02 0.03 0.03 0.03 0.02 Mg 0.98 0.82 0.73 0.53 0.60 0.72 0.76 0.51 0.54 Ca 3.95 3.98 3.98 3.91 3.96 3.98 3.95 3.99 3.80 Sample No. Wairakite Subzone Laumontite-wairakite-bearing rock AW331 ALW02 A P 0 V A P 0 C M SiO2 36.36 37.14 37.47 36.79 36.72 36.43 37.84 36.47 36.04 Al2O3 19.90 20.27 19.79 18.41 17.82 20.56 20.31 19.18 12.31 FeO* 11.82 10.73 11.59 13.96 14.69 11.11 9.70 12.31 24.24 MnO 0.27 0.24 0.30 0.36 0.28 0.05 0.45 0.73 0.09 MgO 2.97 3.75 3.32 2.00 2.56 2.68 4.01 2.76 1.17 CaO 22.29 22.95 22.77 22.25 22.58 22.80 22.70 22.37 21.77 Total 93.61 95.08 95.24 93.77 94.65 93.63 95.01 93.82 95.62 Atomic proportions Si 5.92 5.92 5.99 6.05 5.98 5.92 6.02 5.95 6.01 Al 3.82 3.81 3.73 3.57 3.42 3.94 3.81 3.69 2.42 Fe 1.61 1.43 1.55 1.92 2.00 1.51 1.29 1.68 3.38 Mn 0.04 0.03 0.04 0.05 0.04 0.01 0.06 0.10 0.01 Mg 0.72 0.89 0.79 0.49 0.66 0.65 0.95 0.67 0.29 Ca 3.89 3.92 3.90 3.92 3.94 3.97 3.87 3.91 3.89 FeO* means total iron as FeO. Igneous andmetamorphicpetrology of theHorokanai ophiolite 33 pumpellyite occurs as prisms or needles in vein cite. Of these, the olivine pseudomorph is identi- and amygdule, measuring 0.05-0.2 mm long, up to fied by its shape and the existence of inclusion of 0.5 mm long, and commonly associated with relict spinel. Pleochroism of pumpellyite is also zeolites, albite, calcite and rarely quartz. Pumpel- various, but seems to be related to its mode of lyite in amygdule often forms spherulitic sheaves occurrence; i.e. pale green to green in plagioclase, of acicular crystals with rotary extinction between green to deep green along or around clinopy- crossed nicols. In interstitial matrix, pumpellyite roxene and in olivine pseudomorph, deep green to forms discrete patches accompanied by chlorite, brown in vein, amygdule and interstitial matrix, sphene and Fe-Ti oxide. In most of pumpellyite- and occasionally yellow to gold in interstitial mat- bearing rocks of Zone A, in which pervasive re- rix. crystallization and deformation are feeble, the Assuming that the cation sites are totally fil- precursor minerals of pumpellyite could be ascer- led without vacancy, Coombs et al. (1976) have tained, and the following modes of occurrence are proposed the following general chemical formula observed; replacing plagioclase as veinlets for pumpellyite: W4X2Y4Z,O(20+x)OH(8-x), together with zeolites, albite and rarely quartz, re- where W=Ca, Mn; X=(Mg, Fe2+, Mn) (2-x) placing clinopyroxene as cluster of very fine nee- (Fe3+, Al)x; Y=Al, Fe3+; Z=Si, Al. From reli- dles along cracks or around margins of clinopy- able pumpellyite analyses reviewed by Passaglia & roxene, and replacing olivine pseudomorph as Gottardi (1973), the value of (Mg, Fe2+, Mn) in aggregates associated with chlorite and rarely cal- the X-site ranges from 0.71 to 1.55, that is, the 5.0 4.5 4.0 3.5 3.0 2.5 2.0 1.5 1.0 1.0 1.5 2.0 2.5 3.0 3.5 4.0 0.5 1.0 AI Mg Fig. 21. Compositional variation of Al and Mg against Fe* for pumpellyites. 34 HIDEO ISHIZUKA AI AI 06 80 AL531 AW562 ALW02 Mg 70 . 60 C CA APO 50 MV H v:filling vein A: filling amygdule P:replacing plagioclase o:replacing olivine C:replacing clinopyroxene Mg M:occupying matrix Fe* Fe* Fe* Fig. 22. Compositional variations of Al-Fe*-Mg ratios for pumpellyites occurring in various modes within in- dividual samples. AL531, AW562 and ALW02 represent the laumontite-bearing, wairakite-bearing and laumontite-wairakite-bearingrocks,respectively. filling filling eplacing placing vein amygdule /plagioclase olivine Cpx. matrix Laumontite-bearing rocks :Wairakite-bearing rocks * : Lau.-Wai.-bearing rocks Fe Fig. 23. Compositional variations of Al-Fe*-Mg ratios for pumpellyites occurring in various modes and rock- types. Igneous and metamorphic petrology of the Horokanai ophiolite 35 value of (Fe3+, Al) in the same site ranges from 1.29 to 0.45 per formula unit. This chemical for- CHLORITE mula was used in the present study, and the ato- :ZONEA mic proportions were calculated on the basis of :ZONEB 16.0 cations. O:ZONE C For the Horokanai pumpellyites, we may de- :ZONE D pict several features. The Si value in the Z-site is close to the ideal 6. The Fe3+ value in the Y-site is considerably high, and then Al in the X-site is absent. Such an iron-rich composition may be most characteristic of pumpellyite from zeolite facies metamorphic terrains (Boles & Coombs, 1977; Liou,1979; Evarts & Schiffman, 1983). The X-site is in excess of the ideal figure of 2 while the available Ca and Mn in the W-site cannot fill the ideal value of 4, but the total occupants of the X- 50 and W-sites nearly equal 6. This suggests that th d so s n X-site, may also enter the W-site. On Fe*-Al and Fig. 24. Compositional variation of Al-Fe*-Mg ratio Fe*-Mg diagrams (Fig. 21), Fe* varies antithetical- for chlorites (after Ishizuka, 1985). ly to Al, and Mg tends to decrease with increasing Fe*. Besides these general features, there are pro- Chlorite composition is generally uniform nounced variations in pumpellyite composition within individual samples, except for mixed-layer against the mode of occurrence, as shown in Al- smectite/chlorite clay in Zone A that has the com- Fe*-Al diagrams (Figs. 22 and 23). Within indi- positions varying considerably even within a single vidual samples, the most Al-rich pumpellyite sample; e.g.S SiO2=35-42 wt.%，l Al2O3=12-17 Occurs in plagioclase, while the most Fe*-rich vari- wt.%, FeO*=15-27 wt.%, MgO=10-14 wt.%. ety occurs in matrix. The pumpellyites in other The Fe*/(Fe*+Mg) ratio in chlorite decreases sys- modes have intermediate compositions. Excep- tematically with increasing metamorphic grade, tionally high Fe* contents (as much as 32 wt.%), ranging from 0.43 (Zone A) to 0.27 (Zone D). approaching julgoldite (Fe2+-Fe3+ pumpellyite: Since the FeO*/MgO bulk rock ratio in chlorite- Moore, 1971), were found in pumpellyites with bearing rocks ranges from 0.98 to 1.30 but shows pleochroism ranging from yellow to gold that no sign of systematic variation against metamor- occur in interstitial matrix of the laumontite- phic grade (Table 6), it is suggested that the Fe*/ bearing rocks. The pumpellyite associated with cli- (Fe*+Mg) ratio of chlorite tends to decrease with nopyroxene tends to be rich in Mn. increasing metamorphic grade. Similar composi- Comparing the compositions of pumpellyites tional trends have been described in other meta- occurring in similar mode (Fig. 23), it is found morphic terrains (e.g. Cooper, 1972; Kurata & that the pumpellyites from the laumontite-bearing Banno,1974). Onthe contrary, the SiO2/ rocks are systematically higher in Fe* (thus lower (SiO2+AlzOs) ratio in chlorite is nearly constant in Al) than those from the wairakite-bearing throughout all the mineral zones (0.63±0.03), rocks; the wairakite-bearing rocks, as described which contradicts a tendency that this ratio de- previously, were recrystallized at higher tempera- crease with increasing metamorphic grade as de- tures than the laumontite-bearing ones. These scribed in many metamorphic terrains (cf. Fig. 3 observations indicate the temperature dependence of Maruyama et al., 1983). of the AlzO, content of pumpellyite. 4-4-4. Epidote (Table 10) 4-4-3. Chlorite (Table 9) Epidote is common in Zone B, but sporadic in Chlorite is abundant in Zones A, B and C, but Zone C; the frequency distribution of Xps=Fe3+/ rare in Zone D; the compositional variation in (Fe3+ +Al) is illustrated in Fig. 25. terms of Al-Fe*-Mg against metamorphic grade is In Zone B, epidote is usually zoned with Xps illustrated in Fig. 24. decreasing from the core to the rim. Comparing 9 Table 9. Representative analyses of chlorites from Horokanai Zones A, B, C and D. Sample No. Zone A Zone B Zone C Zone D AC151 AL351 AL331 AL651 AW351 B1501 B1502 B1503 C1501 C2501 C4301 C5501 C6501 C7301 C8501 D1501 SiO2 29.21 29.56 29.30 27.89 28.29 28.45 28.81 29.43 29.40 29.29 29.15 Al2O3 16.55 16.03 16.51 15.57 17.11 16.82 17.10 17.20 18.56 18.05 18.32 17.61 17.21 17.85 17.93 17.99 FeO* 23.77 24.35 24.25 22.41 23.05 22.75 21.62 21.22 19.67 20.85 18.74 18.79 18.19 16.89 16.11 15.89 MnO 0.30 0.29 0.23 0.31 0.28 0.26 0.27 0.35 0.32 0.17 0.38 0.30 0.33 0.31 0.27 0.25 MgO 18.15 17.91 17.95 19.70 18.67 19.51 20.69 20.59 21.51 21.29 22.22 22.82 23.47 23.70 24.52 24.73 HII CaO 0.08 0.09 0.12 0.08 0.07 0.04 0.07 0.10 0.08 0.10 0.12 0.08 0.11 0.09 0.07 0.08 D DEO Total 88.06 88.23 88.36 88.21 88.24 88.03 88.14 88.11 88.03 88.75 88.23 88.41 88.74 88.24 88.19 88.09 ZIHSI Atomic proportions, O=28.0 Si 6.061 6.137 6.070 6.193 5.993 5.9245 5.835 5.875 5.681 5.743 5.750 5.811 5.890 5.873 5.833 5.807 UKA Al 4.049 3.923 4.033 3.772 4.159 4.100 4.143 4.158 4.457 4.320 4.365 4.187 4.061 4.204 4.210 4.225 Fe 4.124 4.228 4.2023 3.8513 3.975 3.934 3.7163.638 3.351 3.540 3.168 3.169 3.045 2.822 2.683 2.648 Mn 0.052 0.051 0.0400.0540.048 0.046 0.0470.060 0.055 0.029 0.066 0.051 0.057 0.053 0.045 0.042 Mg 5.612 5.5425.5426.0335.7376.0136.3376.292 6.530 6.441 6.693 6.860 7.001 7.056 7.277 7.342 Ca 0.0170.0200.0260.0170.0150.0090.0150.022 0.017 0.022 0.026 0.017 0.024 0.019 0.014 0.017 FeO*means total iron as FeO. Igneous andmetamorphic petrology of the Horokanai ophiolite 37 Table 10. Representative analyses of epidotes from Horokanai Zones B and C. SampleNo. Zone B Zone C B1501 B1502 B1503 C1501 C4301 C5501 core-rim core-rim core-rim SiO 37.58 37.95 37.39 38.10 37.74 38.16 38.48 38.39 38.45 TiO2 0.01 0.03 0.05 0.02 0.02 0.03 0.02 0.03 0.02 AlO 22.95 25.46 23.42 26.18 24.02 26.98 28.35 28.92 28.99 FeO,* 13.77 10.52 12.97 9.41 12.41 8.27 6.70 6.05 6.00 MnO 0.15 0.17 0.23 0.15 0.20 0.21 0.10 0.12 0.11 MgO 0.04 0.02 0.07 0.02 0.09 0.12 0.03 0.02 0.03 CaO 23.33 23.59 23.16 23.68 23.41 23.65 23.75 23.85 23.88 Total 97.83 97.74 97.29 97.56 97.89 97.42 97.43 97.38 97.48 Atomicproportions, O=12.5 !S 3.002 2.997 2.996 3.001 2.998 2.998 3.001 2.990 2.991 Ti 0.001 0.002 0.003 0.001 0.001 0.002 0.001 0.002 0.001 A1 2.161 2.370 2.212 2.431 2.249 2.499 2.607 2.655 2.659 Fe 0.828 0.625 0.782 0.558 0.742 0.489 0.393 0.355 0.351 Mn 0.010 0.011 0.015 0.010 0.013 0.014 0.007 0.008 0.007 Mg 0.005 0.002 0.008 0.002 0.011 0.014 0.003 0.002 0.003 Ca 1.997 1.996 1.988 1.999 1.992 1.991 1.985 1.990 1.990 FeO*means total iron asFeO 38 HIDEOISHIZUKA + + 10l + ZONE B EPIDOTE ZONE C 0.1 0.2 0.3 0.4 Fe3+/(Fe3+ + Al) Fig. 25. Frequency distribution of Fe3+/(Fe3+ +Al) ratio for epidotes (after Ishizuka, 1985). Sample An-Content PLAGIOCLASE ZONE No. 10 30 50 70 90 D4401 D4302 D4301 D3301 D3502 D3501 D2503 D2502 D2501 -range- D1501 o C9301 C9501 average C8501 o C7301 C6501 o C C5501 O o C4301 C2501 C1501 B1503 81502 B B1501 AL331 A AL351 Fig. 26. Variation of An-content for plagioclases (after Ishizuka, 1985). Sample No. is arranged from the bot- tom to the top in the ascending order of metamorphic grade. samples from the low-grade part (No. B1501) and increasing metamorphic grade. On this basis, che- the high-grade part (No. Bi503) in Zone B, the mical zoning of Zone B epidotes is interpreted to minimum Xps (i.e. Xps of the rim) is higher in the represent the prograde stage of metamorphism. former (Xps=0.21) than the latter (Xps=0.16). Jilin Late in composition, and has lower Xps (0.08-0.15) than (epidote-actinolite-chlorite-albite-quartz-sphene) that of Zone B. and a similar Fe*/(Fe*+Mg) ratio in chlorite (0.38±0.02). The identified opaque phases include 4-4-5. Plagioclase (Table 11) ilmenite, magnetite and pyrite, but no hematite. Plagioclase is ubiquitous in all the mineral zones; Such a relationship between Xps and metamorphic (Ma) grade is consistent with theobservation of phic grade is illustrated in Fig. 26. SK In Zones A and B, plagioclase is albite, and HG Early mum Xps of the epidote-chlorite-actinolite assemb- metamorphic grade, ranging from 0.9 in the lage is temperature dependent and decreases with lower-grade part of Zone A (No. AL351) to 4.2 in Igneous and metamorphic petrology of the Horokanai ophiolite 39 Table 11. Compositions of plagioclases from Horokanai Zones A, B, C and D. Zone K20 Fe2O3* An-content Number SampleNo. wt. % wt. % JO range average average average analyses AL351 A et al., 2017). In the Jilin area, Wang et al. (2015) determined that 0.9 0.09 0.09 9 AL331 A ly-Middle Paleozoic arc basement, given that the Upper Paleozoic strata 1.0 0.12 0.27 AW351 A 0.8-3.0 1.3 0.13 0.22 12 B1501 B within Upper Permian strata are widely recognized in the South Kita- 2.0 0.08 0.25 15 B1502 B 1.0-4.7 2.3 0.20 0.37 15 B1503 B 1.4-6.3 4.2 0.10 0.10 15 C1501** C 1.8-7.0 4.6 0.07 0.21 49 20.1-24.7 22.3 0.12 0.24 9 C2501** C 1.3-8.5 6.0 0.11 0.37 23 18.9-22.8 21.2 0.13 0.39 6 diate volcanism during the Early to Middle Permian to a mature arc C 1.9-9.5 6.4 0.05 0.24 17 C5501** C 1.8-7.5 5.0 0.06 0.17 21 of the Permian strata of Northeast China-Inner Mongolia, and discuss 19.3 0.09 0.11 9 C6501 C 1.5-5.8 4.0 0.17 0.07 16 C7301** C 1.3-5.0 3.1 0.13 0.43 17 mudstone, volcanic rocks, and volcaniclastic sandstone interbedded 17.3 0.24 0.45 10 C8501** C 1.1-3.8 2.5 0.12 0.31 10 mation consists of lava and volcaniclastic rocks containing rhyolite and 15.0 0.13 0.20 5 C9501** C 0.7-3.7 1.9 0.21 0.11 13 10.1-14.5 12.2 0.20 0.12 7 C9301** C 0.5-2.1 1.2 0.10 0.08 13 8.0-11.3 9.7 0.14 0.09 9 by the input of volcanic detritus in a sedimentary basin, decrease D 9.2-20.3 15.7 0.27 0.29 24 D2501 D 18.9-35.4 23.9 0.12 0.29 18 D2502 D 27.6-41.4 33.8 0.19 0.35 12 D2503 D gradually upward through the stratigraphy (Fig. 11C). On the basis of 38.7 0.07 0.23 9 D3501 D 44.3-50.2 48.8 0.15 0.09 13 temperate/Tethyan species, has been recognized broadly across Inner D age peak at 264 Ma (Capitanian), together with a number of grains with 50.6 0.29 0.17 15 D3301 D Baolidao arc belt, and Hegenshan back-arc belt, respectively. The 62.9 0.22 0.33 11 D4301 D 62.4-69.3 65.3 0.17 0.43 19 D4302 D 75.0-82.5 77.2 0.12 0.39 13 D4401 D 77.6-89.2 85.4 0.20 0.44 19 FezOs* means total iron as Fe2O3. Samples with asterisks (**) contain two-phase plagioclases. 40 HIDEOISHIZUKA the higher-grade part of Zone B (No. B1503). the Horokanai peristerite gap toward the lower In Zone C, both albite and oligoclase An-content with increasing metamorphic grade. (peristerite pair) coexist as discrete grains (Plate This does not support the concept of a solvus and 3G), and a distinct compositional gap can be de- is favourable to a transitional loop, but certainly fined. With increasing metamorphic grade, the more data are needed to substantiate this rela- average An-content in albite first increases from tionship. On the other hand, the calcite-bearing 4.6 (No. C1501) to 6.0 (No. C2501) and then de- samples from Zone C (Nos. C4301 and C6501) creases from 5.0 (No. C5510) to 1.2 (No. C9301), contain albite but are devoid of oligoclase, sug- whereas the average An-content in coexisting oli- gesting that Pco2 may be a factor that plays a role goclase decreases successively from 22.3 (No. in inhibiting oligoclase-producing reaction(Cooper, C1501) to 9.7 (No. C9301). As shown in Table 11 1972). by the relative amounts of albite and oligoclase in In Zone D, one-phase plagioclase occurs. individual samples, increasing metamorphic grade Although a large range of the An-content is en- appears to result in the formation of greater countered for each sample, the average An- amounts of oligoclase in comparison with albite. A preliminary study on a few grains shows that phic grade, ranging up to bytownite. the albite grains are relatively uniform in composi- All the analysed plagioclases are extremely tion, but two types of compositional zoning are low in Fe2O3* and K2O, less than 0.5 and 0.3 detected in the coexisting oligoclase grains; one wt.%, respectively. has a sodic margin, and the other a calcic one, the width of the margin being less than 10 μm in both 4-4-6. Amphibole (Tables 12 and 13) types. Orville (1974) has considered that the Amphibole occurs in Zones B, C and D; the variation of AliV against metamorphic grade is peristerite gap is a two-phase binary loop involv- ing crystallographic transformation in albite-rich illustrated in Fig. 27,in which the double circles composition. However, many descriptions on pointed to by "Table 12" nearly correspond to the natural rocks have shown that the peristerite gap maximum AllV frequency within individual sam- is an asymmetrical solvus delineated by a steep ples, and the plot of (Na+K) and Fe*/(Fe*+Mg) sodic limb and a gentle calcic limb (e.g. Fig. 7 of versus Al'V is shown in Fig. 28. Maruyama et al., 1982) or by a gentle sodic limb In Zone B to the lower-grade part of Zone and a steep calcic limb (Tagiri, 1973). As is clearly C, amphibole is actinolite, and its maximum AlIV shown in Fig. 26, there is a compositional shift of frequency /increases gradually with increasing Sample AIIV AMPHIBOLE No. ZONE 0.5 1.0 1.5. D4401 D4302 D4301 D3301 D3502 D3501 D D2503 D2502 D2501 -range D1501 C9301 C9501 Table 12 C8501 C7301 C6501 C C5501 C4301 C2501 C1501 B1503 B1502 B B1501 Fig. 27. Variation of AlV for amphiboles (after Ishizuka, 1985). Sample No. is arranged from the bottom to the top in the ascending order of metamorphic grade. Igneous andmetamorphic petrology of theHorokanai ophiolite 41 1.0 :ZONE B AMPHIBOLE ●:ZONEC O:ZONED K + 2 0.5 品 0.5 0.5 1.0 1.5 2.0 AIIV represents a coexisting actinolite and hornblende pair; both analyses are listed in Table 12. grade of metamorphism, ranging from 0.10 (No. acterized by a dominant age peak at 416 Ma, subsequent peaks B1501) to 0.50 (No. C2501). and minor Proterozoic zircons. A similar age distribution is recognized In the medium-grade part of Zone C, both hornblende, and a slightly asymmetrical gap re- actinolite and hornblende coexist as discrete acicu- sults. The Fe*/(Fe*+Mg) ratio is commonly lower lar to prismatic crystals (Plate 3H ), and a distinct in actinolite than its coexisting hornblende. In- compositional gap can be defined. With increasing asmuch as the FeO*/MgO bulk rock ratio of two- sandstone (sample NM12-183) in the upper part of the formation has a in actinolite increases from 0.61 (No. C4301) to together, sandstones of the Zhesi Formation are characterized by an source area was a magmatic arc of Ordovician-Early Devonian age (i.e., the Sorayama Formation, and the Daheshen and Dashizhai formations. blende decreases from 1.27 (No. C4301) to 1.04 et al. (1983) regarded the compositional gap be- (No. C6501), and finally both converge at the tween actinolite and hornblende as shifting toward crest about Al'V=0.9. It seems that, with increas- the lower Ai'V content with increasing metamor- 42 HIDEOISHIZUKA Table 12. Representative analyses of Ca-amphiboles from Horokanai Zones B, C and D. Abbreviations: Act=actinolite,Hb=hornblende. SampleNo. Zone B Zone C B1501 B1502 B1503 C1501 C2501 C4301 C5501 Act Act Act Act Act Act-Hb Act-Hb SiO2 54.87 54.09 53.20 52.69 52.03 51.37 45.20 51.02 46.43 TiO2 0.02 0.03 0.04 0.06 0.08 0.11 0.21 0.14 0.25 Al2O3 0.81 1.43 2.01 2.82 3.72 4.82 9.15 5.13 8.77 FeO* 12.81 12.52 12.01 11.64 11.80 11.45 16.02 11.15 14.89 MnO 0.18 0.23 0.20 0.17 0.15 0.17 0.22 0.22 0.20 MgO 15.90 16.53 16.80 17.18 16.43 16.23 12.74 16.65 13.13 CaO 12.96 12.71 12.52 12.63 12.49 12.83 12.45 12.77 12.37 Na2O 0.14 0.30 0.43 0.52 0.63 0.75 1.53 0.81 1.41 K2O 0.08 0.07 0.06 0.05 0.07 0.05 0.27 0.07 0.05 Total 97.77 97.91 97.27 97.76 97.40 97.78 97.79 97.96 97.50 Atomic proportions,O=23.0 Si 7.896 7.775 7.689 7.573 7.513 7.392 6.730 7.327 6.864 Ti 0.002 0.003 0.004 0.006 0.009 0.012 0.024 0.015 0.028 Al 0.137 0.242 0.343 0.478 0.633 0.817 1.606 0.869 1.529 Fe 1.542 1.506 1.452 1.399 1.425 1.378 1.995 1.339 1.841 Mn 0.022 0.028 0.024 0.021 0.018 0.021 0.028 0.027 0.025 Mg 3.409 3.541 3.619 3.680 3.536 3.480 2.827 3.563 2.892 Ca 1.998 1.957 1.939 1.945 1.932 1.978 1.986 1.965 1.960 Na 0.040 0.083 0.120 0.145 0.177 0.209 0.442 0.226 0.403 K 0.014 0.012 0.010 0.009 0.012 0.009 0.052 0.012 0.009 Sample No. Zone C Zone D C6501 C7301 C8501 C9501 C9301 D1501 D2501 Act-Hb Act-Hb Hb Hb Hb Hb Hb SiO2 49.75 47.61 50.37 48.44 49.29 46.95 45.75 46.41 47.24 TiO2 0.18 0.33 0.15 0.41 0.58 0.55 0.53 0.57 0.61 Al2O3 5.67 8.16 5.35 7.04 6.99 9.17 11.10 8.97 6.87 FeO* 10.74 13.92 10.51 12.85 11.30 13.27 14.13 11.13 16.61 MnO 0.24 0.22 0.25 0.21 0.29 0.24 0.23 0.21 0.25 MgO 17.29 13.91 17.40 14.54 15.67 13.68 11.85 16.54 12.40 CaO 12.83 12.51 12.84 12.57 12.45 12.45 12.18 12.53 12.24 Na2O 0.92 1.25 0.87 1.03 1.17 1.59 1.79 1.31 0.99 K2O 0.08 0.07 0.08 0.10 0.09 0.08 0.13 0.04 0.11 Total 97.70 97.98 97.82 97.19 97.83 97.98 97.69 97.71 97.32 Atomic proportions, O=23.0 Si 7.177 6.958 7.240 7.093 7.112 6.853 6.725 6.739 7.042 Ti 0.020 0.036 0.016 0.045 0.063 0.061 0.059 0.062 0.069 Al 0.965 1.406 0.906 1.215 1.189 1.578 1.924 1.536 1.208 Fe 1.296 1.701 1.264 1.573 1.364 1.620 1.737 1.352 2.071 Mn 0.029 0.027 0.030 0.026 0.036 0.030 0.028 0.026 0.031 Mg 3.717 3.030 3.727 3.173 3.369 2.976 2.596 3.579 2.755 Ca 1.983 1.959 1.978 1.972 1.925 1.947 1.918 1.949 1.956 Na 0.257 0.355 0.242 0.292 0.328 0.451 0.511 0.368 0.287 0.014 0.012 0.014 0.019 0.017 0.014 0.025 0.007 0.021 个 Igneous andmetamorphicpetrologyoftheHorokanai ophiolite 43 (Table 12. continued) Sample No. Zone D D2502 D2503 D3501 D3502 D3301 D4301 D4302 D4401 Hb Hb Hb Hb Hb Hb Hb Hb SiO2 47.91 48.53 47.83 46.89 46.00 44.30 44.39 43.80 TiO2 0.74 0.67 0.88 0.95 0.83 1.20 1.47 1.79 Al2O3 6.85 6.71 7.49 7.72 9.91 10.96 11.00 11.83 FeO* 15.21 15.24 14.82 14.53 13.81 12.93 11.62 10.23 MnO 0.25 0.21 0.21 0.22 0.24 0.19 0.25 0.16 MgO 13.51 12.99 12.58 13.36 13.38 13.39 14.30 15.10 CaO 12.28 12.43 12.48 12.44 12.08 12.00 11.94 11.89 Na2O 1.00 0.84 0.94 1.00 1.46 2.01 2.23 2.36 K2O 0.09 0.08 0.03 0.04 0.10 0.04 0.06 0.05 Total 97.84 97.70 97.26 97.15 97.81 97.02 97.26 97.21 Atomic proportions, O=23.0 Si 7.050 7.136 7.058 6.939 6.743 6.551 6.517 6.398 Ti 0.082 0.074 0.098 0.106 0.092 0.133 0.162 0.197 A1 1.188 1.163 1.303 1.347 1.713 1.910 1.904 2.037 Fe 1.872 1.874 1.829 1.798 1.693 1.599 1.426 1.250 Mn 0.031 0.027 0.027 0.028 0.030 0.024 0.031 0.020 Mg 2.963 2.847 2.767 2.946 2.923 2.951 3.129 3.287 Ca 1.936 1.958 1.973 1.972 1.898 1.901 1.878 1.861 Na 0.285 0.240 0.270 0.286 0.416 0.576 0.635 0.669 K 0.018 0.014 0.005 0.007 0.019 0.007 0.011 0.009 FeO* means total iron as FeO. Table 13. Representative analyses of sodatremolites from the Horokanai dunites. Sample No. STD-1 STD-2 STD-3 SiO2 57.45 57.43 57.93 57.36 57.72 TiO2 0.01 0.01 0.01 0.01 0.02 Al2O3 1.77 1.47 1.58 1.31 1.95 Cr2O3 0.77 1.20 0.51 0.83 0.81 FeO* 1.38 1.40 1.50 1.24 1.17 MnO 0.05 0.06 0.05 0.04 0.04 MgO 22.71 22.09 22.87 22.92 22.57 CaO 9.92 9.81 9.89 9.91 9.21 Na2O 4.52 4.64 4.23 4.33 4.63 K20 0.05 0.02 0.03 0.02 0.05 Total 98.63 98.13 98.60 97.97 98.17 Atomic proportions, O=23.0 Si 7.816 7.859 7.866 7.848 7.859 Ti 0.001 0.001 0.001 0.001 0.002 Al 0.284 0.237 0.253 0.211 0.313 Cr 0.083 0.130 0.055 0.090 0.087 Fe 0.157 0.160 0.170 0.142 0.133 Mn 0.006 0.007 0.006 0.005 0.005 Mg 4.606 4.507 4.630 4.675 4.581 Ca 1.446 1.438 1.439 1.453 1.344 Na 1.192 1.231 1.114 1.149 1.222 K 0.009 0.003 0.005 0.003 0.009 FeO* means total iron as FeO 44 HIDEO ISHIZUKA In the medium- to higher-grade part of Zone Ed-Pa D where pyroxene occurs, the Al'v content of hornblende increases again with increasing meta- morphic grade to approach pargasite. 1.5 All amphiboles described above are poor in Na(M4) less than 0.14 and well within the low- pressure field on the Alvi versus Si diagram of Raase (1974). Furthermore, the hornblendes coex- 1.0F Rich isting with sphene or ilmenite show a systematic ·: Horokanai dunites enrichment in TiO2 with increasing metamorphic grade, ranging from 0.21 wt.% (No. C4301) to ±:Buell Park kimberlite 0.5 : Green Knobs kimberlite 1.79 wt.% (No. D4401). This is in harmony with +: Langban metamorphosed Raase's (1974) empirical rule of temperature de- limestones pendence of the TiO2 content. Sodatremolite, which occurs only in the du- 0.5 1.0 1.5 Na nite layers within the harzburgite, has the follow- ing chemical formula: (Nao.6) (Ca1.4Nao.6) (Mg4.6 Fe2+0.2Cro.1Alo.1) (Alo.2Si7.8)O22(OH)2. The com- Fig. 29. Ca-Na plot for sodatremolites in the dunite, positional variation of sodatremolite as illustrated along with those in the Langban meta- morphosed limestones (Deer et al., 1963), in Fig. 29 is mainly due to the substitution of *Ca and in kimberlite diatremes of the Buell for NaNa in the *Ca2MgsSisOz2(OH)2(tremolite)- Park (Aoki et al., 1972) and·the Green Na2CaMgsSisO22(OH)2(richterite)S solid solution Knobs (Smith, 1979) (after Ishizuka, 1980b). series. Such substitution is also recognized in Abbreviations: Tr-Tsch=tremolite-tscher- sodatremolite occurring in metamorphosed lime- makite, Ed-Pa=edenite-pargasite, Rich=- stones (Deer et al., 1963, Table 55) and kimber- richterite, Gl=glaucophane. lites (Aoki et al., 1972; Smith, 1979). 4-4-7. Pyroxene (Table 14) phic grade, and suggested that such a composi- Pyroxene develops in the medium- to high-grade tional shift was attributed to a transitional loop part of Zone D, and in the ultramafic rocks; the associated with a crystallographic transformation compositional variation in the pyroxene quadri- in actinolite-rich composition. The compositional lateral is illustrated in Fig. 30. gap as shown in Fig. 27 shows no sign of such a In the orthopyroxene-free rocks of Zone D, compositional shift. In Sample C7301 which occurs clinopyroxene is salite with Al2O3 of 0.8-1.2 in the higher-grade part of Zone C than Sample wt.%, and shows a systematic decrease of CaO C6501, actinolite coexisting with hornblende has a with increasing metamorphic grade as in the maximum Al1v frequency of 0.79 lower than that amphibolite facies rocks of Central Abukuma of Sample C6501, and then the actinolitic limb of (Shido, 1958) and Broken Hill (Binns, 1962). the gap appears to bend toward the lower AliV These clinopyroxene compositions @ differform content. However, this actinolite is not a discrete those of relict clinopyroxenes persisting in Zone D crystal but a lamella in a host hornblende. Com- as porphyroclasts; these relict clinopyroxenes are paring only discrete crystals of coexisting actinolite augites with Al2O3 of 4.5-5.1 wt.%. and hornblende, the compositional gap delineated In the orthopyroxene-bearing rocks of Zone in the present study, especially the successive D, clinopyroxene is salite to augite with AlzO, of change of actinolitic limb compositions, favours a 1.0-2.1 wt.%, and orthopyroxene is hypersthene solvus over a transitional loop. with CaO less than 0.7 wt.% and AlzO; of 0.5-1.3 In the higher-grade part of Zone C to the wt. %. The partition coefficient, Kd=(Fe*/Mg)Opx/ lower-grade part of Zone D, amphibole is horn- (Fe*/Mg)Cpx between coexisting pyroxenes, ranges blende, but its Al'V content varies considerably from 1.6 to 1.8, which is within the limits of meta- from one rock to another. Such variation may be morphic pyroxene pairs (Kd=1.5-2.0: Kretz, due to lack of buffered assemblages in the analy- 1963). The Mori-Green (1978) thermometer gives sed samples except for Sample D1501 which in- equilibration temperatures of 660°C±40°C for cludes chlorite. such Kd values, if the FezOs content is assumed to Sample No. Zone D Ultramaficrocks D3501 D3502 D3301 D4301 D4302 D4401 U3010 U3021 Cpx Cpx Cpx Cpx-Opx Cpx-Opx Cpx-Opx Cpx-Opx Cpx-Opx and SiO2 52.01 52.08 52.12 52.95 52.12 52.81 52.73 53.31 53.95 53.82 56.72 53.92 56.82 met TiO2 0.08 0.07 0.07 0.17 0.04 0.15 0.04 0.16 0.04 0.01 0.01 0.01 0.01 Al2O3 1.14 0.89 1.17 2.04 1.24 1.53 0.99 1.06 0.56 0.87 0.61 0.84 0.51 Cr2O3 0.03 0.03 0.02 0.02 0.02 0.04 0.02 0.03 0.02 0.21 0.10 0.24 0.10 FeO* 9.78 9.58 9.12 8.32 23.26 7.72 21.37 7.14 19.28 2.26 7.02 2.20 6.52 petr MnO 0.37 0.35 0.39 0.29 0.36 0.20 0.35 0.34 0.35 0.11 0.22 0.10 0.22 13.12 13.75 14.13 22.13 14.75 23.84 15.15 25.10 19.05 rology MgO 12.45 34.55 19.16 34.79 CaO 23.80 23.57 22.99 22.65 0.69 22.43 0.62 22.26 0.68 23.15 0.56 23.20 0.76 Na2O 0.34 0.10 0 0.32 0.26 0.24 0.37 0.01 0.01 0.40 0.01 0.09 0.03 0.03 the Total 99.98 99.95 99.87 99.94 99.87 99.97 99.97 99.85 99.99 99.57 99.82 99.77 99.76 Atomicproportons,O=6.0 Si 1.962 1.962 1.957 1.940 1.954 1.961 1.957 1.976 1.979 1.960 1.969 1.960 1.971 Ti 0.002 0.002 0.002 0.005 0.001 0.004 0.001 0.004 0.001 0.000 0.000 0.000 0.000 Al 0.051 0.040 0.052 0.090 0.055 0.067 0.043 0.046 0.024 0.037 0.025 0.036 0.021 Cr 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.001 0.006 0.003 0.007 0.003 Fe 0.309 0.302 0.286 0.260 0.729 0.240 0.663 0.221 0.592 0.069 0.204 0.067 0.189 0.012 0.011 0.012 0.009 0.011 0.006 0.011 Mn 0.011 0.011 0.004 0.006 0.003 0.006 Mg 0.700 0.737 0.770 0.786 1.237 0.816 1.319 0.837 1.372 1.034 1.787 1.038 1.798 Ca 0.962 0.952 0.925 0.906 0.028 0.893 0.025 0.884 0.027 0.904 0.021 0.904 0.028 Na 0.024 0.019 0.018 0.027 0.001 0.025 0.001 0.029 0.001 0.007 0.002 0.007 0.002 FeO*meanstotal iron asFeO. 46 HIDEOISHIZUKA 10 15 50 A Hd 30 (A) D350 D3502 D3301 D4302 45 D4401 D4301 U3010 PYROXENE Fs (B) 25 30 35 (B) ： 75 70 65 60 Tie line represents a coexisting clino- and orthopyroxene pair; both analyses are listed in Table 14. be negligible. (Mg+Fe2+) ratio ranges from 0.53 to 0.72 in In the ultramafic rocks, clinopyroxene is pyroxenite, from 0.44 to 0.51 in dunite, and from diopside with AlzO3 less than 0.9 wt.%, and 0.45 to 0.58 in harzburgite,while the Cr/ orthopyroxene is bronzite to enstatite with CaO (Cr+Al+Fe3+) ratio ranges from 0.10 to 0.63 in and Al2O, less than 0.8 and 0.6 wt.%, respective- pyroxenite, from 0.78 to 0.88 in dunite, and from ly. The Kd ranges from 1.6 to 1.7, and gives 0.67 to 0.77 in harzburgite. The TiO2 content of equilibration temperatures, by the above thermo- spinel is less than 0.3 wt.%. The olivine-spinel meter, of 680°C±20°C. Orthopyroxene occurring thermometer of Roeder et al. (1979) gives ascoarse-grainedporphyroclast in thecumulate equilibration temperatures of 608-717°C for rocks has the Mg/(Mg+Fe*) ratio of 0.89-0.91, pyroxenite, 598-753°C for dunite, and 598-729°℃ which is more magnesian as compared with coex- for harzburgite. isting cumulus olivine (Mg/(Mg+Fe*)=0.82-0.85) and clinopyroxene (Mg/(Mg+Fe*)=0.75-0.80) and rescmbles the orthopyroxene composition from 5.METAMORPHIC CONDITIONS the basal ultramafic rocks described above. 5-1. Prograde Metamorphism The temperatures of the zeolite facies estimated 4-4-8. Olivine and Spinel (Table 15) from burial metamorphic terrains and active Olivine and spinel are common constituents in the geothermal regions range from 100°C to 300°℃ ultramafic rocks. The forsterite content of olivine, (for a review, see Zen & Thompson, 1974). Low- Fo=Mg/(Mg+Fe*)=0.91-0.93, is relatively con- temperature zeolites such as phillipsite and clinop- stant without reference to parental rock-types. tilolite, which are common Na-zeolites reported in The NiO content of olivine is 0.35 wt.% on aver- deep-sea sediments and stable at less than 50°℃ age. Spinel spans a wide range of composition, (e.g. Boles, 1977), do not appear to occur in although the Fe3+/(Cr+Al+Fe3+) ratio is small re- Horokanai Zone A. The experimental studies of gardless of parental rock-types; e.g. the Mg/ zeolites showed that, at Phuid= Ptotal=2kb, the sta- Igneous andmetamorphicpetrology of theHorokanai ophiolite 47 Table 15. Chemical compositions of olivines and spinels, and temperature estimates for Horokanai ultramafic rocks. Sample No. Olivine Spinel T** Mg* Mg* Cr* A1* Fe* (°C) Pyroxenite P-3010 0.918 0.719 0.096 0.804 0.100 717 P-3021 0.918 0.626 0.355 0.580 0.065 665 P-4010 0.930 0.534 0.628 0.338 0.034 603 Dunite D-5010 0.918 0.507 0.832 0.089 0.079 753 D-5021 0.918 0.494 0.853 0.130 0.017 710 D-5031 0.926 0.503 0.847 0.060 0.093 723 D-3010 0.927 0.440 0.878 0.064 0.058 630 D-3021 0.929 0.461 0.781 0.178 0.041 598 Harzburgite H-5010 0.921 0.453 0.766 0.230 0.004 598 H-5021 0.923 0.474 0.714 0.254 0.032 603 H-5032 0.928 0.584 0.711 0.251 0.038 729 H-3010 0.926 0.528 0.671 0.300 0.029 633 H-3021 0.931 0.523 0.681 0.290 0.029 610 Mg*=Mg/(Mg+Fe2+). Cr*, Al* and Fe*=Cr/(Cr+Al+Fe3+), Al(Cr+Al+Fe3+) and Fe3+/(Cr+Al+Fe3+), respectively. (a) i ia p jo p o q sind=** spinel formula. bility temperatures are less than 170°C for stilbite Zealand and Taiwan. (Liou, 1971c), between 170°C to 280°C for Nitsch (1971) demonstrated experimentally laumontite (Liou, 1971a), more than 280°C for that, at Pruid=Ptota=2kb, the prehnite-pumpel- wairakite, and less than 200°C for analcime + lyite-chlorite-quartz assemblage is stable up to quartz (Liou, 1971b; Thompson, 1971). Mixed- 345°C±20°C, whereas the chemically equivalent layer smectite/chlorite clay with widely varying assemblage actinolite-chlorite-epidote-quartz exists composition as observed in Horokanai Zone A above 350°C; the former assemblage occurs in the can occur at temperatures as high as 200°C (Hoff- highest-grade part of Horokanai Zone A, and the man & Hower, 1979). Few experimental data are latter one is critical to Horokanai Zone B. These available on the stability of iron-rich pumpellyite results, especially the high-temperature stability in the zeolite facies rocks, even though iron-free limit of pumpellyite, have been pursued by the re- Mg-Al pumpellyite has been synthesized cent experimental study of Schiffman &Liou temperatures as low as 250°C but only at (1980). Also,Schiffman & Liou (1980) showed Puid=Ptotal more than 7kb (Schiffman & Liou, that the prehnite-tremolite-chlorite-albite-quartz 1980). For naturally occurring iron-rich pumpel- assemblage as found in the highest-grade part of lyite, Boles & Coombs (1977) suggested 190°C for Horokanai Zone A occupies a relatively small P-T its first appearance in the zeolite facies sandstone field of Pnuid=Ptotal=2-5kb and 325-375°C. of New Zealand, and Liou (1979) ）estimated The mineral assemblage transitional from the temperatures for the laumontite + iron-rich pum- greenschist to amphibolite facies, i.e. actinolite + pellyite assemblage in the East Taiwan ophiolite calcic-plagioclase + chlorite as observed in Horo- to be at 150-250°C; the iron content of the Horo- kanai Zone C, is stable within the temperature in- kanai pumpellyite is similar to those from New terval between 475-550°C at Pnuid=Ptotal=2kb and 48 HIDEO ISHIZUKA at the oxygen fugacity of the QFM buffer (Liou et tion may have been unrelated to the formation of al., 1974), or between 370-420°℃ as given by the antigorite occurring in one sample of the Horoka- carbon isotope and calcite-dolomite thermometers nai ultramafic rocks as “relict" partially replaced (Maruyama et al., 1982, 1983). The discrepancy by chrysotile or lizardite, because antigorite is between the temperatures of Liou et al. (1974) stable at temperatures in excess of 350°℃ and as and Maruyama et al. (1982, 1983) may be due to high as 500°C (Evans et al., 1976; Coleman, 1977). the equality of Ppuid=Ptotal and the lower oxygen It is, however, ambiguous to determine the forma- fugacity in experiments compared to natural tive stage of "relict" antigorite. mineral parageneses, of which the effect of in- creasing oxygen fugacity has been experimentally 5-3. Metasomatism determined to expand the P-T field of the transi- Sodatremolite has been described from meta- tion zone (Apted & Liou, 1983). morphosed limestones (Deer et al., 1963, Table The temperatures of the highest-grade part of 55) and kimberlites (Aoki et al., 1972; Smith, Zone D (hornblende-granulite facies) are 620- 1979), but the Horokanai sodatremolite is the first 700°C as estimated on the Fe-Mg partitioning of report of its occurrence in ultramafic rocks of coexisting pyroxenes calibrated by Mori & Green ophiolites. (1978). The equilibration temperatures of the Aoki et al. (1972) suggested that sodatremo- underlying ultramafic rocks are 660-700°C also on lite was a primary phase in kimberlitic magma the Fe-Mg partitioning of coexisting pyroxenes formed under upper mantle conditions. Hariya & and 600-750°C on the olivine-spinel thermometer Terada (1973) showed by high-pressure synthetic of Roeder et al. (1979), similar to those calculated experiments that tremolite(50)-richterite(50) solid for the overlying hornblende-granulite facies rocks solution has an extremely wide stability field, of Zone D. being stable up to 700°C at 40kb and up to 900°℃ There are few constraints available for the at 34kb. If the Horokanai sodatremolite was an quantitative estimates of the metamorphic press- igneous phase, the host dunite should have been ures. However, the absence of the typical albite- in equilibrium with kimberlitic magma, because epidote amphibolite facies assemblage, the pre- this amphibole has never been known as an sence of the prehnite-actinolite-chlorite assemb- igneous phase in any igneous rocks except kimber- lage in the highest-grade part of Zone A, the pre- lites. The Horokanai ultramafic rocks was in sence of the albite-oligoclase-actinolite-horn- equilibrium with tholeitic magma but not with blende-chlorite-epidote assemblage in Zone C, kimberlitic magma. It is geologically unnatural to consider that only a part of the sodatremolite- as well as Na(M4), all suggest low-pressure condi- bearing dunite coexisted with kimberlitic magma tions (Shido,1958; Miyashiro, 1973; Raase, 1974; while the majority of the ultramafic rocks coex- Brown, 1977; Maruyama et al., 1983; Liou et isted with tholeitic magma. It is, therefore, un- al.,1985). likely that the ：Horokanai sodatremolite was Accordingly, it is suggested that the Horoka- formed as an igneous phase. nai ophiolite was progressively metamorphosed in The second possibility for the origin of the the temperature range of 100-750°C, and its meta- Horokanai sodatremolite is a subsolidus reaction morphic facies series, ranging from the zeolite, in closed-system. As no sodium-bearing phase greenschist, amphibolite to granulite facies, be- occurs in the host dunite, we cannot specify the longs to the low-pressure type. source of sodium. In the Horokanai ultramafic rocks, clinopyroxene occurs only in the pyroxenite 5-2. Serpentinization and gabbroic layers, both of which are rare rock- The serpentinization of the Horokanai ultramafic types. Amphibole except for sodatremolite is ab- rocks, characterized by the chrysotile-lizardite sent in the ultramafic rocks. For these reasons, assemblage, occurred at temperatures less than the present author thinks it improbable to produce 350℃C (Evans et al., 1976; Coleman, 1977). As the sodatremolite by a closed-system subsolidus there is no textural evidence showing recrystalliza- reaction in ultramafic compositions. tion of serpentine minerals, the stage of serpenti- The remaining possibility is a local meta- nization may have postdated the equilibration somatic process which increased the Na/Al ratio in stage of the ultramafic rocks at the granulite the host dunite, even though the source of sodium facies. Also, such low-temperature serpentiniza- and the process of metasomatism still remain un- Igneous and metamorphic petrology of the Horokanai ophiolite 49 solved. In this connection, it may be noted that a short vertical sequence with the facies bound- the local metasomatic process seems to have oper- aries subparallel to the ophiolite pseudostratigra- ated in another ultramafic rocks of the Kamuiko- phy. However, the maximum metamorphic grade tan zone, as was described by Nagata (1982) on attained is in an upper part of the gabbros, below the Iwanai-dake ultramafic rocks. In this ultrama- which the metamorphic effect dies off rapidly. fic rocks, metasomatism gave rise to formation of This boundary between metamorphosed and un- pargasite, monticellite and perovskite, although metamorphosed gabbros is extremely irregular, the elements added to the host rock are different and the deeper penetration of the metamorphic from the Horokanai ultramafic rocks. However, in effect is concentrated along fractures, veins and any case, both ultramafic rocks have common fea- shear zones (e.g. the ophiolite of Chile). tures, that is, the sodium-rich amphibole was (3): Igneous textures are commonly preserved formed and the basicity of the host rocks in- in the zeolite to greenschist facies rocks, but in creased by metasomatism. the amphibolite facies rocks they are largely mod- ified to metamorphic textures such as nematoblas- tic, granoblastic or gneissose textures (e.g. the 6. ORIGIN OF OPHIOLITE METAMORPHISM ophiolites of Alps and Borneo). 6-1. Ophiolite metamorphism; Nature and (4): Petrochemically, the low-grade rocks of Comparison the zeolite to greenschist facies are more or less The metamorphosed ophiolites of the low-pressure accompanied by element migration.A Although type have been described from many localities there are some differences in the degree or extent such as the Mediterranean (Gass & Smewing, of element migration among the ophiolites, the 1973; Spooner, 1974; Spooner & Fyfe, 1973; following general trends of element migration can Spooner et al., 1974; Pamic et al., 1973), Alps be noted during the low-grade metamorphism; (a) (Mevel et al., 1978), Oman (Coleman, 1977), the increases of SiO2 and NazO, (b) the decrease Newfoundland (Coish, 1977), Chile (DeWit & of CaO,(c) the slight decreases of MgO and Stern, 1976; Stern et al., 1976; Elthon & Stern, FeO*, (d) the immobility of TiO2, ZrOz, Cr2O3, 1978; Stern & Elthon, 1979), Taiwan (Liou, 1979; NiO, Y2O3, P2Os and REE, (e) the increase or Liou & Ernst, 1979), Borneo (Hutchison, 1975, decrease of KzO. The vertical change from the 1978), California Coast Range (Evarts and Schiff- man, 1983), and Hidaka Western Zone (Miyashita depleted high-grade metabasites is also reported in et al.， 1980; Miyashita, 1981, 1983). The rela- the Mediterranean and Chilean ophiolites. tionships between the lithology and the metamor- These observations led Coleman (1977) and phic facies series of the ophiolites referred to others to deduce that the grade of the low- above are summarized in Fig. 31, in which the pressure ophiolite metamorphism is lower than the Yakuno ophiolite (Ishiwatari, 1985) is also shown low amphibolite facies, and that the ophiolite as an example of the medium-pressure type of metamorphism does not take place below the up- metamorphism. From Fig. 31, we may depict per part of the gabbroic rocks. several features of ophiolite metamorphism. The metamorphic nature of the Horokanai (1): The mineral facies series generally in- ophiolite is broadly similar to that inferred for the cludes the zeolite, greenschist and amphibolite ophiolite metamorphism given above. However, facies. Particularly, the zeolite facies rocks of the the granulite facies as recognized in the highest- Taiwan and Coast Range ophiolites sometimes grade part of the Horokanai Zone D and its contain iron-rich pumpellyite. The prehnite- underlying ultramafic rocks is generally absent in pumpellyite facies is absent in most ophiolites, but other ophiolites. The orthopyroxene-bearing rocks the Newfoundland ophiolite contains it as the from the lower part of the mafic rocks in other lowest-grade facies instead of the zeolite facies, ophiolites have been generally regardedas and the Coast Range ophiolite includes it as a magmatic (commonly called norite or two- transitional facies between the zeolite and green- pyroxene gabbro), even though few chemical data schist facies. The prehnite-actinolite facies and the are available on them. greenschist-amphibolite transitional facies are re- However, at least the following three ported from the Taiwan ophiolite. ophiolites include the high-grade facies such as the (2): The metamorphic grade increases down- high amphibolite and granulite facies: (a) In the wards from the zeolite to amphibolite facies within ophiolite of Borneo, Hutchison (1975) described S OPHIOLITE METAMORPHISM New- California Troodos Oman Coast Chile Borneo Yakuno Taiwan foundland Range :PP 亿 RB PP PP GS UM UM UM C ? (km) (km) DEO NZIHSI Horokanai EXPLANATION LITHOLOGY METAMORPHICFACIES SedimentaryRocks? Zeolite Facies, Prehnite-Pumpellyite F. Basaltic R. B Greenschist F. GS Doleritic R. EA Epidote-Amphibolite F. T Transition F. Gabbroic R. At'+ Amphibolite F. (f:fault) (km) UM Unmetamorphosed Ultramafic R. GX Granulite F Fig. 31. Nature of ophiolite metamorphism. Data are from the papers listed in text. Igneous and metamorphicpetrology of the Horokanai ophiolite 51 the granulite facies in the lower part of the gab- (R/V) such as ATLANTIS, VEMA, PILLSBURY broic rocks, but the same author (Hutchison, and JEAN CHARCOT. They were mainly col- 1978) has later reinterpreted them as igneous lected from the lower walls of oceanic ridges or cumulates and ophitic gabbro after analysing the transverse fracture zones, and include metabasalts, minerals by microprobe. Hutchison's (1978) two- metadolerites and metagabbros. Furthermore, the pyroxene data, however, give equilibration dredged samples from the fracture zones are com- temperatures of 650-830°C by using the Mori- monly accompanied by serpentinites, and appear Green (1978) thermometer. (b) In the ophiolite of to be inclusions therein. Probably, the serpenti- the Hidaka Western Zone, Miyashita et al. (1980) nites, which were formed by hydration of oceanic described the sapphirine-bearing rocks, and inter- upper mantle peridotites, were emplaced in the preted it as the metamorphosed troctolite recrys- solid state along the fracture zones, tearing apart tallized under the conditions of the high amphibo- the metamorphic rocks recrystallized at some lite to low granulite facies. (c) In the Yakuno depth, and bringing them up to the surface. It is, ophiolite, Ishiwatari (1985) assigned the medium- therefore, most likely that the dredged samples pressure granulite facies to the cumulates,and cover the oceanic materials to a considerable ascribed it to a thermal regime attained beneath a depth, but they provide no direct information on relatively thick oceanic crust like that of the Black metamorphic sequence of oceanic crust. Sea. (2): The drilled or dredged oceanic metaba- salts, metadolerites and metagabbros generally re- 6-2. Ocean-Floor Metamorphism; Nature and tain igneous structure and texture, but sometimes Comparison exhibit metamorphic structure or texture to be Ever since Matthews et al. (1965) reported meta- greenschist or amphibolite. Secondary minerals morphosed basalts dredged on the crest of the develop in veins, fractures and interstitial mat- Carlsberg Ridge, the mid-ocean ridge in the rices, and sometimes occur partially to pervasively northwestern Indian Ocean, widespread occurr- replacing igneous minerals. ences of metamorphic rocks are increasingly being (3): Most of the oceanic metabasalts belong recognized from the ocean-floor over the world. to the zeolite to greenschist facies, while most Particularly, petrology on metamorphic rocks dril- oceanic metagabbros are in the greenschist to led or dredged from ocean-floor have thrown new amphibolite facies. The reported metamorphic light on the metamorphic processes taking place minerals of each metamorphic facies are as fol- near divergent plate boundaries (Melson & van lows: Andel, 1966; Cann & Funnell, 1967; Cann, 1969; The zeolite facies; natrolite-mesolite-scolecite Ploshko et al., 1970; Miyashiro et al., 1971; Au- series, analcime, stilbite, heulandite, mixed layer mento et al., 1971; Bonatti et al., 1 1975;Helm- chlorite/smectite. staedt, 1977; Ito & Anderson, 1983; Honnorez et The greenschist facies; actinolite, hornblende, al., 1984). Miyashiro et al. (1971) called it “ocean- tremolite, albite, chlorite, epidote, quartz, sphene, floor metamorphism". talc, magnetite, nontronite. By reviewing these literatures, we may depict The amphibolite facies; hornblende, actinolite, several features for ocean-floor metamorphism as calcic-plagioclase, chlorite, epidote, biotite, summarized in Fig. 32. quartz, apatite, magnetite, sphene, leucoxene. (1): The drilled samples have been mainly Very rarely, the prehnite-pumpellyite facies meta- provided by the Deep Sea Drilling Project that basalts has been reported, .including prehnite, has been, using the drilling vessel (D/V) GLO- pumpellyite, chlorite, epidote, white mica, albite, MAR CHALLENGER, performed in the ocean- quartz, and calcite. However, the detailed mineral floor over the world. They include mainly metaba- assemblages of the oceanic samples, especially of salts with minor metadolerites. Since the deepest the low-grade rocks, are not specified in most drilling hole established by 1985 is of 1075.5 m literatures except for Mevel (1981) who has de- basement (Anderson et al., 1982), most drilled scribed the prehnite-pumpellyite facies metabasites samples represent the shallow materials of oceanic dredged from the Vema fracture zone of the Paci- crust. However, they supply direct information on fic Ocean and recognized the following mineral assemblages: metamorphic sequence of oceanic crust. On the other hand, the dredged samples have been albite-chlorite-pumpellyite-prehnite-white mica- obtained during many cruises of research vessels quartz-calcite, S OCEAN-FLOOR METAMORPHISM INDIAN OCEAN PACIFIC ATLANTIC OCEAN Carlsberg Ridge OCEAN Mid-Atlantic Ridge 5°S 6N E.P.Rise 45N 30&24'N 22'N N.0-0L PP QB B B EO ISF ZIH 10 10 (km) EXPLANATION (km) LITHOLOGY METAMORPHIC FACIES Sedimentary Rocks Zeolite Facies Basaltic R. PP Prehnite-Pumpellyite F. Doleritic R. Greenschist F. Transition F. Gabbroic R. A+++ Amphibolite F. Uitramafic Granulite R. Fig. 32. Nature of ocean-floor metamorphism. Data are from the papers listed in text. Igneous and metamorphic petrology of the Horokanai ophiolite 53 albite-chlorite-epidote-prehnite-quartz-calcite, ferrous ratio generally increases as compared with albite-chlorite-pumpellyite-quartz-white mica- the basaltic precursors. In contrast, the amphibo- calcite, lite facies metadolerites and metagabbros show no In this respect, recently I have had a chance to sign of element migration. study the metamorphic rocks dredged from the The metamorphic nature of ocean-floor meta- Pacific Ocean by courtesy of Professor A. morphism reviewed above, especially its facies Miyashiro and National Science Museum, and series and original rock association, is similar to observed the following mineral assemblages: that of the Horokanai ophiolite. However, there is Thezeolitefacies one important difference; i.c. thc tcxture of thc laumontite-analcime-pumpellyite-chlorite high-grade rocks is completely metamorphic in the laumontite-thomsonite-chlorite-albite Horokanai ophiolite while the oceanic high-grade The greenschist facies samples usually retain igneous texture. albite-chlorite-actinolite albite-chlorite-actinolite-epidote 6-3. Discussion The transitional greenschist to amphibolite facies Although the origin of ophiolites has been con- albite-oligoclase-actinolite-chlorite-epidote troversial, as shown in the first chapter, the Horo- albite-actinolite-hornblende-chlorite kanai ophiolite has several lines of evidence (e.g. albite-oligoclase-actinolitic hornblende-chlorite- lithology, geochemistry and mineralogy) for its epidote origin at oceanic ridge rather than at island arc or The amphibolite facies oceanic island. Furthermore, similar metamorphic calcic plagioclase-hornblende nature of the Horokanai ophiolite to that of the calcic plagioclase-hornblende-clinopyroxene ocean-floor metamorphism suggests that the origin calcic plagioclase-hornblende-clinopyroxene- of its metamorphism must be discussed in relation orthopyroxene to the origin of ocean-floor metamorphism. The zeolite and greenschist facies assemblages are One of the most important features of the restricted to the metabasalts and metadolerites Horokanai ophiolite and ocean-floor metamorph- while the amphibolite facies ones develop only in ism is an unusually high geothermal gradient. the metagabbros. However, the transitional facies Although no complete sequence is available to assemblages are found both in the metadolerites estimate the original thickness of the Horokanai and metagabbros. Most importantly, the studied ophiolite, an order of 100°C/km is inferred for its samples, even of the amphibolite facies, contain geothermal gradient. Such a geothermal gradient relict clinopyroxene and sometimes plagioclase. is consistent with other ophiolite metamorphism Metamorphic minerals usually develop replacing (Spooner & Fyfe, 1973; Coleman 1977) except the relict clinopyroxenes along margins or crack, or Yakuno ophiolite which is supposed to have had occupying interstitial phases, or filling veins or the geothermal gradient of 30-45°C/km (Ishiwa- fractures: Plate 6 shows photomicrographs of stu- tari, 1985). In oceanic realms, such a high- died oceanic samples with mineral assemblage of geothermal gradient has been measured only each mineral facies described above. around oceanic ridges; i.e. the geothermal gra- (4): Petrochemically, the element migration dient has been directly measured on Legs 69, 70 of oceanic metabasites occurs during ocean-floor and 83 of DSDP, 201 km south of the Costa Rica metamorphism, but the degree of element migra- Ridge, to be 70-120°C/km (Anderson et al., 1982). tion seems to be different from sample to sample, Another important feature of the Horokanai and to be dependent upon the metamorphic ophiolite and ocean-floor metamorphism is the grade. The zeolite facies metabasalts show great downward increase of metamorphic grade with increases of NazO and HzO. In the greenschist thermal structure subparallel to lithology (Figs. 31 facies metabasalts and metadolerites, two types of and 32). the mineral assemblages are recognized: ：the Volcanism, which is the most possible source chlorite-rich and the epidote-rich assemblages. of heat for metamorphism, occurs in restricted re- The chlorite-rich assemblages gain significant gions in ocean-floor, and is most active in oceanic quantities of MgO and loss CaO, but they tend to ridges. It follows that the tectonic setting responsi- retain the original ferric/ferrous ratio. The altera- ble for such high-geothermal gradient and thermal tion to the epidote-rich assemblages results in only structure as described above is an oceanic ridge, small migration in CaO and MgO, but the ferric/ and that an episode of the Horokanai ophiolite 54 HIDEOISHIZUKA and ocean-floor metamorphism can takesplace the ocean floor, are unknown, and further there is only near or beneath oceanic ridges. an uncertainty whether such a magma chamber is The igneous process of oceanic ridges has long lived (steady-state) or ephemeral. These are been modeled by many petrologists, and it seems important keys to evaluate thermal structure of to be widely accepted that a magma chamber ex- oceanic ridges. ists beneath oceanic ridges, involving low-pressure In this respect, the recent thermal models of fractional crystallization, and producing a three- the formation of the oceanic crust (Sleep, 1975; layered structure of oceanic crust: an upper unit Kusznir & Bott, 1976; Kusznir, 1980) have shown of basaltic pillow lavas and dikes, a middle unit of that (1) fo1 oceanic ridges with a half-spreading massive gabbro formed by gradual freezing of the rate less than 1 cm/year, the latent heat of basalt margins of the chamber, and a lower unit of in the top 5 km of the oceanic crust is insufficient layered (cumulate) gabbro formed on the base of to maintain a steady-state magma chamber, and the chamber (e.g. Cann, 1974; Hekinian et al., (2) the size of magma chamber is smaller for slow- 1976; Bryan & Moore, 1977; Stakes et al., 1984). spreading ridges than for fast-spreading ridges; This is called the “infinite onion" model (Fig. 33). e.g. the magma chamber has a half width of 10 However, the structural features of such a magma km for a half-spreading rate of 6 cm/year and a chamber, e.g. the size, shape and depth beneath half width of 1.5 km for a half-spreading rate of (km) 5 cm/year SEA LEVEL HYDROTHERMAL. 1 2 BASALT 300 DOLERITE 600 4 800 6 1000 GABBRO 1100 8 1200 CUMULATE 10 PERIDOTITE 10 (Wx 10 20 1 cm/year 0 S.L. HYDROTHERMAL 乙 BASALT 100 4 300 DOLERITE 6 GABBRO 600 8 CUMULATE 800 10 PERIDOTITE 1100 1000 Fig. 33. Thermal structure at and around oceanic ridge with half spreading rate of 5 cm/year and 1 cm/year. Igneous process for faster spreading ridge, called the “infinite onion" model, are after Bryan & Moore (1977), and that for slower spreading ridge, called the “infinite leek" model, are after Nisbet & Fowler (1978). Isotherms (°C) are drawn by the thermal model of Sleep (1975), Kuszir & Bott (1976) and Kuszir (1980). The penetration of hydrothermal cel is shown by the curved arrows near the sea floor. Igneous and metamorphic petrology of the Horokanai ophiolite 55 1.5 cm/year. This means that a steady-state crustal 1973; Hart, 1970, 1973). Since there are many magma chamber can exist in fast-spreading ridges lines of evidence for interaction between seawater such as the East Pacific Rise with a half-spreading and basaltic rocks, e.g. element migration, recrys- rate of 6-8 cm/year (Goslin et al., 1972), but it is tallization of hydrous minerals such as clays and highly questionable whether any steady-state crus- zeolites, and enrichment of 18O, such a circulation tal magma chamber can exist, or it is small or would be attained within basaltic rocks. Ex- ephemeral, in slow-spreading ridges such as the perimental investigations also support such an in- Mid-Atlantic Ridge with a half-spreading rate of teraction to occur possibly (Hajash, 1975; Bischoff 0.7-1.5 cm/year (Pitman & Talwani, 1972). For & Dickson, 1975). There is, however, little evi- these reasons, Nisbet & Fowler (1978) has pro- dence available for the circulation within gabbroic posed the alternative model for slow-spreading rocks. This leads many authors to deduce that the ridges, called the “infinite leek" model, in which hydrothermal circulation does not reach such a rising melt segregations from the mantle first depth. reach the base of the pre-existing oceanic crust, On the basis of petrologic and geophysical and then the melt fill cracks to propagate up from models described above, we may illustrate a ther- the base of the crust at the ridge axis, freezing to mal structure at and around oceanic ridges, as form new crust (dikes), and if it cracks to reach shown in Fig. 33. The geothermal gradient is high the surface of the oceanic crust, the melt is ex- near a ridge axis but decreases with distance away truded to form pillow lavas (Fig. 33). from the axis. Lateral movement of newly formed These thermal models are consistent with oceanic crust will subject rocks at each pseudos- seismic experiments: seismic data obtained from tratigraphic level to a progressively decreasing the Mid-Atlantic Ridges have indicated that a geothermal gradient. Consequently, the shallower magma chamber defined by crustal low-velocity basalts cool quickly while the metamorphism with- zones does not at present exist (Fowler, 1976, in the gabbros occurs gradually in a declining ther- 1978; Steinmetz et al., 1977), in contrast the crus- mal gradient. This process is a kind of autometa- tal low-velocity zones which have been proposed morphism. Furthermore, the geothermal gradient to exist along the ridge axis of the East Pacific depends upon the spreading rate, that is, for the Rise are generally assumed to be magma cham- slower-spreading rate it decreases more rapidly bers with half width of approximately 10 km away from the ridge axis. Thus, it is likely that for (Orcutt et al., 1975; Rosendahl et al., 1976). The the faster-spreading rate the oceanic materials depth of low-velocity zone, i.e. the depth of mag- have enough time to give rise to mineral re- ma chamber, is about 5 km beneath sea level (i.e. equilibrium during their divergence from the ridge about 2.5 km beneath sea floor) in the East Paci- axis. If the degree of recrystallization depends fic Rise (Reid et al., 1977). upon the duration of metamorphism, it is there- These facts suggest that the nature of magma fore likely that the metamorphic rocks formed at chamber and possibly related thermal structure of the faster-spreading ridge are more metamorphic Oceanic ridges are highly dependent upon spread- in texture and mineral chemistry. ing rate. The relation between the thermal struc- As described previously, the degree of recrys- ture and spreading rate of oceanic ridges has been tallization is higher in the Horokanai ophiolite mathematically modeled by Sleep (1976), Kusznir than in dredged metamorphic rocks from the Mid- & Bott (1976) and Kusznir (1980), in which the Atlantic Ridge. The rate of half-spreading is 1.1 geothermal gradients in fast-spreading ridges are cm/yearfor the Mid-Atlantic Ridge in the less steep than those found in slow-spreading FAMOUS region (Macdonald, 1977). As to the ridges (Fig. 33). This is consistent with the fact half-spreading rate for the Horokanai ophiolite, that the observed heat flow decreases more rapid- the following two points suggest that the spreading ly away from oceanic ridges with slow-spreading rate was probably high. First, the Horokanai rates (Le Pichon & Langseth, 1969), that is, the ophiolite lacks the sheeted dike complex. As is heat flow over fast-spreading ridges is greater seen in Fig. 33, for slower-spreading rates the than over slow-spreading ridges at all distances. proportion of intrusive crust is greater than for On the other hand, hydrothermal circulation faster-spreading rates, that is, for faster-spreading of seawater may be another factor operating rates the oceanic crust is predominantly plutonic ocean-floor metamorphism, especially its low- while for slower-spreading rates it is completely grade part (Scott & Hajash, 1976; Thompson, intrusive. This is because at slow-spreading rates 56 HIDEOISHIZUKA the heat supplied by injected materials is con- tion of Japan (Nos. 56740351/Ishizuka, 59740418/ ducted and advected away at a suficiently high Ishizuka, 60740441/Ishizuka). rate to cool the injected materials immediately be- low the solidus temperature (Kusznir, 1980). Second, the recent study of Engebretson et al. APPENDIX. FIELD TRIP (1984) have indicated that the Jurassic oceanic The field trip to follow the “continuous"” sequence plate accompanied by the Kamuikotan geology is of the Horokanai ophiolite, i.e. from the top the Izanagi Plate with the very high rate of half- radiolarian chert to the basal harzburgite, through spreading (10-15 cm/year). Most probably thc dif- the basaltic and gabbroic rocks, needs hard walk, ference of the degree of recrystallization between and then the outcrops of representative rock-types the Horokanai ophiolite and dredged metamor- of easy access were chosen here "un- phic rocks from the Mid-Atlantic Ridge may be continuously". The places visited are shown in explained by the difference of the rate of half- Fig. A-1. spreading. STOP-1:AT THE HAKUSAN ELEMENTARY ACKNOWLEDGEMENTS SCHOOL I would gratefully acknowledge my supervisor, Layered cumulus rocks (boulders) Professor Shohei Banno, for his continuous help, These boulders are precious garden rocks of this advice and encouragement throughout this study, school,so “DO NOT TOUCH THE HAM- without which this thesis would not have been MER!"; they are easily found on the river bed of completed. I am also indebted to the pioneers of the Minami-Jyunisen River, and their outcrops are the research of the Kamuikotan geotectonics, located at the upperstream of the same river. The Messrs. Nobuo Gouchi and Masayuki Imaizumi, most characteristic of the rocks is rhythmic layer- with whom I enjoyed frank discussions during the ing structure defined by the alternation of olivine- course of the study. gabbro, gabbro, dunite and plagioclase-peridotite My sincere thanks are due to Professor Aki- with minor wehrlite and pyroxenite. The thickness ho Miyashiro and National Science Museum who of each layer varies widely, ranging from 1 cm or donated me their ocean-floor samples. less to more than 10 cm. Under the microscope, 1 sincerely thank Professor Masao Yamasaki, cumulus texture is clear, although it has been Professor Seiji Hashimoto, and Professor Takashi generally modified by recrystallization. Occa- Suzuki for their valuable suggestions and con- sionally, mineral grading is distinct of a olivine- tinuous encouragements. I am also grateful to gabbro layer, mainly defined by decrease of oli- many colleagues of the Kanazawa University and vine and increase of plagioclase from bottom to the Hokkaido University for their help in the field top in a layer. and laboratory works and for their daily discus- sions, especially Mr. Masami Nakagawa and D1. Yoichi Motoyoshi for their patiences in teaching STOP-2: AT THE INUSHIBETSU RIVER-1 XRF and AAS techniques. Close-packed pillow lavas with associated dolerite I would like to thank Dr. Makoto Okamura dikes for his identification of radiolarians, and Dr. The basaltic pillow lavas commonly occur as close- Yutaka Takigami for giving me his unpublished packed type, and their outcrops appear bulky, Ar4-Ar9 age data of metamorphic hornblendes rounded and hummocky (Fig. A-2). Individual pil- from the Horokanai ophiolite. lows are highly variable both in shape and size, D1. Hiroaki Sato, Dr. Shigenori Maruyama, although they are generally ellipsoidal in cross sec- Dr. Toshio Higashino, Dr. Akira Ishiwatari, Dr. tion with the maximum and minimum diameters Sumio Miyashita, Dr. Shin-ichi Yoshikura and Dr. ranging from 30 to 80 cm and from 20 to 50 cm, Toshisuke Kawasaki gave me constructive critic- respectively. A glassy rind (chilled margin) of ism and stimulating discussion at various stages of several millimeters in thickness is formed at the the study, which were always a source of inspira- pillow margin. On the pillow surface, the glassy tion and encouragement. part is traversed by a polygonal network of shr- This study was financed in part by Grant-in- inkage cracks (chilled cracks), which extends in- Aid for Scientific Research of Ministry of Educa- wards forming radial columnar joints. Vesicles or Igneous andmetamorphicpetrologyof theHorokanai ophiolite 57 45°E 45°N Kamuikotan Zone 40°N EX P LANATION NeogeneVolcanicRocks CretaceousYezoGroup KamuikotanMetamorphic Rocks HOROKANAI- OPHIOLITE (km） Radiolarian Chert Basaltic Rocks hyaloclastite massive lava flow d .0 banded amphibolite Gabbroic Rocks massive amphibolite Cumulate Rocks Ultramafic Rocks b dunite nses foliated harzburgite massive harzburgite HOROKANAI (serpentinite) Mixed-Up Zone) Thrust Fault (barbs on upper plate) Fault Lithological Contact Fig. A-1.Geological map of the Horokanai ophiolite (after Ishizuka, 1980a, 1985), in which the places (Nos. 1-10) visited are shown. The thin arrow in the inset points to the locality of the Horokanai ophiolite. HIDEOISHIZUKA 58 10m illowlava doleritedike amphibolite faultbreccia Fig.A-2.A sketch map along the Inushibetsu River. 10 zW| >SE Fig. A-3.A sketch figure showing far-distance view of the dike-like body of micro-gabbro within the host harz- burgite along the Inushibetsu River. Igneous and metamorphic petrology of the Horokanai ophiolite 59 amygdules are rare in the glassy pillow margin, overlying ultramafic rocks by the low-angle fault, but present sporadically in the crystalline pillow extending northwest to southeast with about 250 core. The dolerite dike develops within the pillow m width and dipping 50°-70° northeast (Fig. A-1). sequence, reaching maximum width of about 5 m, This zone seems to represent one of a tectonic of which the grain size increases from the contact melange zone, in which numerous angular to sub- with the host rocks (pillow lavas) to the inner angular blocks were mixed up (Figs. A-4A and part. The pillow lavas and dolerite dikes common- -4B). The blocks, ranging in size from 50 cm or ly retained their primary igneous texture; inter- less to more than 10 m, are derived from various granular (pillow core), variolitic (pillow margin) rock-types of the Horokanai ophiolite, but do not and subophitic to ophitic (dolerite dike) textures. include the members of the Kamuikotan high- However, they contain abundant secondary miner- pressure metamorphic rocks and Yezo Group. als including zeolites, chlorite, pumpellyite, albite, The formation of this zone is thought to be closely calcite, sphene, Fe-Ti oxide and rarely quartz, related with ophiolite emplacement onto the being strongly indicative of low-grade metamorph- Kamuikotan metamorphics. ism. Such secondary minerals occur as partially or totally replacing primary igneous phases or as wholly filling radiating or concentric fractures and STOP-5: AT THE ENTRANCE OF THE GOSEN veins. RIVER Hyaloclastite (one of the basaltic member) The hyaloclastite are tuff breccia-like rocks having STOP-3: AT THE INUSHIBETSU RIVER-2 many basaltic fragments embedded in a cogenetic Basal harzburgite with some micro-gabbro tuffaceous matrix (Fig. A-5). As a whole, they As a whole, the Horokanai ultramafic rocks have underlie the pillow sequence, but some also been intensely serpentinized, and no fresh perido- occur intercalated into the pillow sequence. The tite remains. However, on the basis of relict fragments are derived largely from pillow basalts mineral assemblages in less serpentinized parts, with minor doleritic rocks, of which both the the parental peridotite types can be identified as shape and size vary from place to place, ranging mainly harzburgite with minor dunite and pyroxe- from angular to subangular or subrounded in form nite. The less serpentinized harzburgite with pro- and from 1 cm or less to more than 10 cm in dia- togranular texture is composed of olivine and meter. Occasionally, the hyaloclastites containing orthopyroxene with minor spinel, but clinopy- numerous fine-grained shell-like glassy pillow frag- roxene is very rare. In more serpentinized rocks, ments are found. The hyaloclastites observed at orthopyroxene is almost entirely altered to bastite, this stop were metamorphosed in the zeolite to and olivine-remnants of serpentinization forms lower-greenschist facies, and exhibit weak folia- fine-rounded grains as floating in a network of tion. magnetite and serpentine minerals. X-ray diffrac- tion shows that serpentine minerals are mainly chrysotile and lizardite. The micro-gabbro, occur- STOP-6: AT THE DOWNSTREAM OF THE ring as dike-like or lenticular bodies within the GOSEN RIVER host harzburgite (Fig. A-3), consists of plagioc- Downwards protruding shape of close-packed lase, clinopyroxene and brown hornblende, but pillows plagioclase is altered to more sodic plagioclase Generally, it is very difficult to obtain the precise with a trace of epidote, and clinopyroxene and attitude of the pillow sequence in the field, be- hornblende are partially replaced by chlorite and/ cause the pillow pile tends to lack other than or blue-green hornblende. It follows that the crude stratification. However, a tendency for pil- micro-gabbro was more or less metamorphosed in lows to flatten in the place of deposition results in the epidote-amphibolite to amphibolite facies. an approximate “bedding". In this respect, it is noteworthy that some pillows are balloon-shaped with circular upper surfaces and prominent tails STOP-4: AT THE EASTERN SIDE OF THE plastically yprotruding downwards,1 fitting re- INUSHIBETSU DAM entrant spaces of underlying pillow surfaces (Fig. "Mixed-Up Zone"-tectonic melange zone (?) A-6). Such “smooth-curved" tops (convex up- So-called “Mixed-Up Zone" is in contact with the wards and tails downwards) provide a criterion to HIDEOISHIZUKA 09 (A) Fig.B amphibolite >tuff zmassive basalt MIXED dolerite UP 7 hyaloclastite R!Y 7pillow(?)basalt ZO rocks 150 m （B） NW >SE 10m massiveamphibolite foliatedamphibolite banded amphibolite ultramatic rock (harzburgite?) with micro-gabbro e o dex, g sso s () ss e () d 't ushibetsu River. Igneous and metamorphic petrology of the Horokanai ophiolite 61 NW SE 3m RIVER GOSEN Fig.A-5.A sketch figure showing far-distance view of the hyaloclastite at the entrance of the Gosen River STOP-7: AT THE MIDDLESTREAM OF THE GOSEN RIVER-1 Radiolarian chert and Yezo Group The total volume of the chert member is extreme- ly small (less than 0.1 per cent) in relation to the apparent exposure of the Horokanai ophiolite. BOTTOM This chert member can be divided into three 50cm types; (1) massive chert, (2) bedded chert, and (3) basaltic fragment-bearing chert. Radiolarians are abundant in the type (1) and (2) cherts, but less Fig. A-6.A sketch figure of close-packed pillows common in the type (3) chert. The red chert showing downwards protruding shape at the observed here belongs to the type (1) chert with downstream of the Gosen River. some basaltic materials, and yields the following radiolarians; Parvicingula hsui, Mirifusus mediodi- latatus, Praeconocaryomma magnimamma, Trico- locapsa sp. and Zhamoidellum sp., indicating the differentiate top and bottom of the pillow sequ- concurrent range of Tithonian. The Yezo Group is ence.Also a tendency for vesicles to concentrate characterized by the alternation of sandstone and in the upper part of the pillow body is useful to mudstone, of which the mudstone also yields the do this, but the amount of vesicles in the Horoka- following radiolarians; Holocryptocanium astiensis, nai pillows is so small to determine the younging H. barbui group and Novixitus sp., indicating the direction of the pillow sequence. On the other concurrent range from Upper Albian to Cenoma- hand, the interpillow reddish matrix made of nian. The existence of the time gap from Titho- basaltic and pelagic materials are found between nian to Upper Albian allows the interpretation of pillows, from which several radiolarians were the geologic relation between the ophiolite and yielded. the Yezo Group as unconformity or fault. 62 HIDEOISHIZUKA STOP-8: AT THE MIDDLESTREAM OF THE GOSEN RIVER-2 Banded amphibolite (one of the basaltic member) The rocks have strong metamorphic features with granoblastic texture; especially the banded struc- ture definedbythe alternation of leucocratic (plagioclase-rich)and melanocratic (amphibole- rich) parallel bands is well developed, each band measuring 1 mm or less to more than 10mm in width.The constituent minerals include horn- blende and plagioclase with minor sphene,Fe-Ti oxide （mainly ilmenite)and rarely apatite，to which epidote and chlorite may be added in some Du:dunite cases. The original rocks of these amphibolites is Hz:harzburgite obscure,because of strong recrystallization. However,it is locally observed in the rocks with weakly-developed banded structure that igneous clinopyroxene partially replaced by hornblende Flg. A-7.A sketch figure of a ultramafic rock (boul- der) showing layering structure of harzbur- shows blastintergranular texture with plagioclase, gite and dunite at the upperstream of the suggesing that the original rocks are largely of Gosen River. basaltic rocks. STOP-9: AT THE UPPERSTREAM OF THE acterized by layering structure. At this stop, such GOSEN RIVER-1 a structure appears as convex (harzburgite) and Massive amphibolite (gabbroic member) concave (dunite) surfaces resulting from the dis- This type of amphibolites is massive, but has tinctive degree of weathering, each layer being pa- sometimes gneissose structure.Under the micro- rallel to each other;especially it is typically scope,it displays granoblastic texture,and occa- observed on weathered boulders on the river bed sionally trace igneous clinopyroxene (augite) is (Fig.A-7). 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MIYASHIRO, A., SHIDO, F. and EwING, M., 1971: Meta- Mem. Geol. Soc. Japan, 6, 61-74. morphism in the Mid-Atlantic Ridge near 24° and MATTHEWs, D.H., VINE, F.J. and CANN, J.R., 1965: 30°N. Phil. Trans. Roy. Soc. London, A-268, 589- Geology of an area of the Carlsberg Ridge, Indian 603. Ocean. Geol. Soc. Amer. Bull., 76, 675-682. MiYAsHITA, S., 1981: Ophiolite succession and meta- McBIRNEy, A.R. and W1LLiAMs, H., 1969: Geology and morphism of the Western Zone of the Hidaka petrology of the Galapagos islands. Geol. Soc. metamorphic belt, Hokkaido. In “Tectonics of Amer. Mem., 118, pp.197. Paired Metamorphic Belts (ed. HARA, 1.)", Pp.25- Melson, W.G., Thompson, G. and ANDEL, T.H., 30, Hiroshima: Tanishi Print Kikaku. 1968: Volcanism and metamorphism in the Mid- MiYAsHrTA, S., 1983: Reconstruction of the ophiolite Atlantic Ridge, 22°N latitude. Jour. Geophys. Res., succession in the western zone of the Hidaka meta- 73, 5925-5941. morphic belt, Hokkaido. Jour. Geol. Soc. 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Radiolarian Symp., BIBEE, L.D., HusSONG, D.M. and SuTTON, G.H., tectonic setting (Arai, 1992). In Fig. 7, the compositions 1976: Evolution of oceanic crust, part I: A physical ORCUTr, J.A., KENNETT, B.L.M. and DoRMAN, L.M., model of the East Pacific Rise crest derived from 1975: Structure of the East Pacific Rise from an seismic refraction data. Jour. Geophys. Res., 81, ocean bottom seismometer survey. Geophys. Jour. 5294-5304. Roy. Astron. Soc.,45, 305-320. SATO, H., AOKI, K., OKAMOTO, K. and FUJITA, B., OrviLLE, P.M., 1974: The peristerite gap as an equilib- 1979: Petrology and chemistry of basaltic rocks rium between ordered albite and disordered pla- from Hole 396B, IPOD/DSDP Leg 46. Initial Rep. gioclase solid solution.Bull. Soc. fr. Mineral. Crys- DSDP, 46, 115-142. tal., 97, 386-419. SCHIFFMAN, P. and LIOU, J.G., 1980: Synthesis and sta- PAGE, B.M., 1981: The Southern Coast Ranges. In “The bility relations of Mg-Al pumpellyite. Jour. 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PEARCE, J.A., 1975: Basalt geochemistry used to investi- SEKI, Y. 1971: Some physical properties of analcime- gate past tectonic environments on Cyprus. Tecto- wairakite solid solutions. Jour. Geol. Soc. Japan, Igneous and metamorphic petrology of the Horokanai ophiolite 69 77, 1-8. Geophys. Res., 89, 6995-7028. SHIBAKUsA,H., 1974: Glaucophane schists in the STARK, J.T., 1963: Petrology of the volcanic rocks of Kamuikotan metamorphic belt of the Horokanai Guam. U.S. Geol. Surv. Prof. Paper, 403-C, pp.32. area, central Hokkaido. Jour. Geol. Soc. Japan, 80, STEINMETZ, L., WHITMARSH, R. and MOREIRA, V., 341-353, (in Japanese with English abstract). 1977: Upper mantle structure beneath the Mid- SHIBAKUSA, H., GOUCH1, N. and IMAIZUMI, M., 1977: Atlantic Ridge north of the Azores based on which is similar to HIMU and the Mino-Tamba rocks, observations of compressional waves. Geophys. in the Kamuikotan metamorphic rocks near Asahi- Jour. Roy. Astron. Soc., 50, 353-380. kawa, Hokkaido. Jour. Geol. Soc. Japan, 83, 301- STERN, C.R., DEW1r, M.J. and LAWRENCE, J.R., 1976: 303. Igneous and metamorphic processes associated with SHIDo, F., 1958: Plutonic and metamorphic rocks of the the formation of Chilean ophiolites and their im- Nakoso and Iritono district in the central Abukuma plication for ocean floor metamorphism, seismic Plateau. Jour. Fac. Sci., Univ. Tokyo, Sec. Il, 11, layering and magnetism. Jour. Geophys. Res., 81, 132-217. 4370-4380. SHIDo, F., MIYAsHIRO, A. and EwING, M., 1971: Crys- STERN, C. and ELTHON, D., 1979: Vertical variations in tallization of abyssal tholeiites. Contrib. Mineral. the effects of hydrothermal metamorphism in Chi- Petrol., 31, 251-266. lean ophiolites: their implications for ocean floor SHIRAK1, K., 1971: Metamorphic basement rocks of Yap metamorphism. Tectonophysics, 55, 179-213. Islands, western Pacific: possible oceanic crust be- SuRDAM, R.C., 1973: Low-grade metamorphism of tuf- neath an island arc. Earth Planet. Sic. Lett., 13, faceous rocks in the Karmutsen Group, Vancouver 167-174. Island, British Columbia. Geol. Soc. Amer. Bull., SHIRAki, K., YusA, Y., KuRODa, N. and IsHIoKa, K., 84, 1911-1922. highly dependant on the selection of the composition Suzuki, J. and Suzuk1, Y., 1958: Petrological study of Mariana island arc. Jour. Geol. Soc. Japan, 83, 49- picrite (Ichiyama and Ishiwatari, 2005; Ichiyama, unpublished data). 57, (in Japanese with English abstract). kaido, Japan. Jour. Fac. Sci., Hokkaido Univ., Ser. SIGURDsSON, N. and SCHILLING, J.G., 1976: Spinels in IV, 10, 349-446. gram, Zr content is changed by variable degrees of TAGIR1, M., 1973: Metamorphic rocks of the Hitachi dis- ence. Earth Planet. Sci. Lett., 29, 7-20. trict in the southern Abukuma Plateau, Japan. Sci. SIGVALDAsON, G.E., 1968: Structure and products of Rep. Tohoku Univ., Ser. Ill, 12, 1-67. subaquatic volcanoes in Iceland. Contrib. Mineral. TAGIRi, M., 1977: Fe-Mg partition and miscibility gap Petrol., 18, 1-16. between coexisting calcic amphiboles from the SLEEp, N.H., 1975: Formation of oceanic crust: some southern Abukuma Plateau, Japan. Contrib. Miner- thermal constraints. Jour. Geophys. Res., 80, 4037- al. Petrol., 62, 271-281. 4042. Tarney, J., SaunDers, A.D. and WEAVER, S.D., SMITH, D., 1979: Hydrous minerals and carbonates in 1977: Geochemistry of volcanic rocks from the peridotite inclusions from the Green Knobs and Fig. 7. Cr-spinel compositions of the Tamba ferropicrite and Mino Buell Park kimberlitic diatremes on the Colorade region. In “Island Arcs, Deep Sea Trenches and Plateau. In "The mantle Sample: Inclusions in Kim- Back-Arc Basin (eds. TalwanI, M. and berlites and Other Volcanics (eds. BoYD, F.R. and PITman, W.C.)", Pp.367-378, WashinGTon, D.C.: MEYER, H.O.A.)", Pp.345-356, Proc. 2nd Inter. Amer. Geophys. Union. Kimberlite Conf., 2, Amer. Geophys. Union. TAYLOR, S.R., ANNETTE, C.C., GRAHAM, A.L. and SpOONER, E.T.C., 1974: Sub-sea-floor metamorphism, BLACK, D.H., 1969: Trace element abundances in heat and mass transfer: an additional comment. andesites. Contrib. Mineral. Petrol., 23, 1-26. Contrib. Mineral. Petrol., 45, 169-173. TAZAKI, K., 1964: Alkali amphibole-bearing metamor- SPOONER, E.T.C., BECKINSALE, R.D., FYFE, W.S. and olivine accumulation or subtraction, while Nb/Zr ratio SMEwING, J.D., 1974: O'8 enriched ophiolitic meta- ern part of Asahikawa, central Hokkaido. Jour. basic rocks from E.Liguria (Italy), Pindos (Greece), Assoc. Geol. Collab. Japan, 71, 8-17, (in Japanese and Troodos (Cyprus). Contrib. Mineral. Petrol., with English abstract). 47, 41-62. TEMPLEMAN, J.A., 1979: Chromian spinels in Leg 49 SPOONER, E.T.C. and FYFE, W.S., 1973: Sub-sea-floor basalts: a preliminary chemical study. Initial Rep. metamorphism, heat and mass transfer. Contrib. DSDP,49,745-748. Mineral.Petrol.,42,287-304. THOMPSON, A.B., 1971: : Analcime-albite equilibria at STAKES, D.S., SHERVAIS, J.W. . and HoPsON, C.A., low temperatures. Amer. Jour. Sic., 271, 79-92. 1984: The volcanic-tectonic cycle of the FAMOUS THOMPsON, G., 1973a: A geochemical study of the low- and AMAR valleys, Mid-Atlantic Ridge (36°47'N): temperature interaction of sea-water and oceanic Evidence from basalt glass and phenocryst composi- igneous rocks. EOS Amer. Geophys. Union Trans., tional variations for a steady state magma chamber 54,1015-1019. beneath the valley midsections, AMAR 3. Jour. THOMPsON, G., 1973b: Trace-element distribution in 70 HIDEOISHIZUKA fractionated oceanic rocks, 2. Gabbros and related rocks. Chem. Geol., 12, 99-111. THOMPSON, G., SHIDO, F. and MIYASHIRO, A., 1972: Trace-element distributions in fractionated oceanic basalts. Chem. Geol., 9, 89-97. WaGer, L.R., Brown, G.M. and Wadsworth, W.J., 1960: Types of igneous cumulates. Jour. Petrol., 1, 73-85. WAGER, L.R. and MrrCHELL, R.L., 1953: Trace ele- ments in a suite of Hawaian lavas. Geochim. Cos- mochim.Acta.,3, 217-223. WaTANABE, J., 1965: Studies ofthe Horokanai amphibolite massif, Kamuikotan structural belt, central Hokkaido. Jour. Geol. Soc. Japan, 71, 120- 137, 193-214, 281-296, (in Japanese with English abstract). WATANABE, T., 1982: Idonnappu Formation in the Kamuikotan belt as constituent of a Mesozoic sub- duction wedge. In “Tectonics of Paired Metamor- phic Belts (ed. HARA, I.)", pp.37-41, Hiroshima: Tanishi Print Kikaku. WHITE, A.J.R., JAKES, P. and CHRISTIE, D.M., 1971: Composition of greenstones and the hypothesis of sea-floor spreading in the Archaean. Spec. Publs. Geol. Soc. Aust., 3, 47-57. WooD, D.A., 1978: Major and trace element variations in the Tertiary lavas of eastern Iceland and their significance with respect to the Iceland geochemical anomaly. Jour. Petrol., 19, 39-436. Wood, D.A., Varet, J., Bougault, H., Corre, O., JoRoN, J.L., TreuIl, M., BIzouarD, H., Nor- RY, M.J., HAWKESWORTH, C.J. and RoD- DICK, J.C., 1979: The petrology, geochemistry, and mineralogy of North Atlantic basalts: a discussion based on IPOD Leg 49. Initial Rep. DSDP, 49, 597-656,885-902. ZEN, E-an. and THOMPsON, A.B., 1974: Low-grade re- gional metamorphism: mineral equilibrium rela- tions. Ann. Rev. Earth Planet. Sci., 2, 179-212. Plates 1-6 Explanation of plate 1 A: Occurrence of chert showing bedded structure composed of chert and thin shale bed. B: Photomicrograph of chert showing many radiolarians. The width of plate is approximately 2.2 mm. One nicol. C: Far-distance view of close-packed pillow lavas. D: Pilw lavas showing curved tops; convex upwards (up side on the plate) and tails downwards (down side on the plate). E: Occurrence of hyaloclastite containing many pillow fragments. F: Photomicrograph of glassy material-bearing hyaloclastite, in which glassy materials are now altered to palagonite or chlorite. The width of plate is approximately 2.2 mm. One nicol. G: Occurrence of tuff layer (TUFF) in hyaloclastite (HYALO). H: Occurrence of massive lava flow (MLF) in hyaloclastite (HYALO). Plate1 Explanationofpliate2 A: Occurrence of banded amphibolite composed of leucocratic (plagioclase-rich) and melanocratic (hornblende-rich) bands. B: Occurrence of massive amphibolite withgneissose structure. C: Occurrence of dolerite dike in massive amphibolite. D: Occurrence of plagiogranite dike in banded amphibolite. E:1 Layered structure of cumulaterock. F:( Close-up view of cumulate rock; the convex and concave surfaces are rich in plagioclase and olivine, respectively. G: Layered structure of ultramafic rock defined by the alternation of dunite (DU) and harzburgite (HZ). H: Occurrence of gabbro layer in ultramafic rock; the width of plate is about 50 m. Plate2 Explanation of plate 3 I'I Ade s aed jo yim o 'e mod uz u wu on io sudon yo ydeoud mm. One nicol. B: Photomicrograph of holocrystalline core in Zone A pillow lava. The width of plate approximately 2.2 mm. One nicol. C: Photomicrograph of interpillow matrix in Zone A pillow lava. The width of plate is approximately 2.2 mm. One nicol. D: Photomicrograph of hyaloclastite in Zone B, in which acicular crystals are actinolites. The width of plate is approx- imately 2.2 mm. Crossed nicols. E: Photomicrograph of Zone C hyaloclastite. The width of plate is approximately 2.2 mm. Crossed nicols. F: Photomicrograph of Zone C banded amphibolite. The width of plate is approximately 2.2 mm. One nicol. G: Photomicrograph of coexisting albite (Ab) and oligoclase (Og) in Zone C hyaloclastite. The width of plate is appro- ximately 1.1 mm. One nicol. H: Photomicrograph of coexisting actinolite (At) and hornblende (Hb) in Zone C hyaloclastite. The width of plate is approximately 1.1 mm. One nicol. Plate 3 Explanation of plate 4 imately 2.2 mm. One nicol. C: Photomicrograph of orthopyroxene (Op)-clinopyroxene (Cp)-bearing hornblende granulite in Zone D. The width of plate is approximately 2.2 mm. One nicol. D: Photomicrograph of relict clinopyroxene (RCp) in Zone D massive amphibolite, partially replaced by hornblende. The width of plate is approximately 1.1 mm. One nicol. E: Photomicrograph of adcumulate texture in cumulate rock. The width of plate is approximately 2.2 mm. Crossed nicols. mm. Crossed nicols. ss w 7 xde si ad jo im u yoi oenn ui anxn rend jo yeiroud nicols. approximately 2.2 mm. One nicol. Plate 4 Explanation of plate 5 A: Photomicrograph of pumpellyite in Zone A rock, filling vein. The width of plate is approximately 2.2 mm. One nicol. s h si d no d Crossednicols. C: Photomicrograph of pumpellyite in Zone A rock, replacing plagioclase. The width of plate is approximately 1.1 mm. One nicol. D: Photomicrograph of pumpellyite in Zone A rock, replacing olivine. The width of plate is approximately 1.1 mm. Onenicol. E: Photomicrograph of pumpelyite in Zone A rock, replacing clinopyroxene; the arrows point to typical habit of pum- pellyite. The width of plate is approximately 1.1 mm. One nicol. The width of plate is approximately 2.2 mm. One nicol. Plate5 Explanation of plate 6 A: Photomicrograph of dredged zeolite facies metabasalt from the Mid-Atlantic Ridge, showing occurrence of thomso- nite (Th). The width of plate is approximately 2.2 mm. Crossed nicols. B: Photomicrograph of dredged greenschist facies metadolerite from the Mid-Atlantic Ridge, showing occurrences of epidote (Ep), actinolite (At) and albite (Ab). The width of plate is approximately 2.2 mm. Crossed nicols. C: Photomicrograph of dredged greenschist-amphibolite transition facies metabasite from the Mid-Atlantic Ridge, Crossednicols. showing occurrence of aggregate (Agg) composed of coexisting albite and oligoclase, in which relict clinopyroxene (RCp) is partially replaced by hornblende. The width of plate is approximately 2.2 mm. Crossed nicols. E:F Photomicrograph of dredged amphibolite facies metagabbro from the Mid-Atlantic Ridge, composed of hornblende (Hb), calcic plagioclase (Pl) and clinopyroxene (Cp). The width of plate is approximately 2.2 mm. One nicol. F: Photomicrograph of dredged granulite facies metagabbro from the Mid-Atlantic Ridge, showing the replacement of relict clinopyroxene (RCp) by aggregate composed of hornblende, calcic plagioclase, clinopyroxene and orthopy- roxene.The width of plate is approximately 2.2 mm. One nicol. Plate 6
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Ishizuka (1987) - Igneous and metamorphoc petrology of the Horokanai Ophiolite.txt
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Geochemical Journal, Vol. 20, pp. 143 to 151, 1986
NOTE
Geochemical implication of the
of saline spring waterslithium content
in Japan
NOBUKI TAKAMATSU', MASAYUKI IMAHASHII, KYOKO KAMIMURA1
and MAKOTO TsuTSUMI2
Department of Chemistry, Toho University, Miyama, Funabashi, Chiba 274',
and Institute for Thermal Spring Research, Okayama University,
Misasa, Tottori 682-022, Japan
(Received September 27,1985: Accepted May 6, 1986)
Lithium as well as major chemical constituents, and stable hydrogen and oxygen isotopic ratios in 30
saline spring waters (Cl>5,000 mg/Q) in Japan were determined to clarify the genesis of the waters. The
relationship between log (Na/Li) and the temperature estimated by geothermometers revealed that the
lithium content of saline spring waters are not usually controlled at the present-day temperatures of the
hydrothermal systems. Taking into account the values of the deviation coefficient of lithium to seawater,
CLi = (Li/Cl) sample/(Li/Cl) seawater, and the regions from which the spring waters were collected, the
spring waters can be classified into the following four groups: Group 1, representing saline waters from
coastal line (CLi = 1.2 22), Group 2 those from greentuff region (CLi = 17 110), Group 3 those from
Osaka Basin (CLi = 160 220), and Group 4 those from the outer part of Median Tectonic Line (CLi =
250 440). The genetic features of the respective groups are discussed in detail.
INTRODUCTION
Some volcanic Na-Cl type or highly saline
waters are enriched in lithium. This fact has
led us to some ideas for the origin of lithium.
White (1957) stated that high lithium concen
tration of volcanic Na-Cl type waters can be
reasonably explained by concentration of
lithium in residual magma and the subsequent
transportation as soluble halides in a dense vapor
phase. On the other hand, Ellis and Mahon
(1964) suggested that high lithium concentra
tion in thermal waters does not necessarily
indicate their magmatic origin. Furthermore,
Takamatsu et al. (1983) showed from the dis
solution experiments of Li minerals in salt
solutions that some Japanese non-volcanic Na-Cl
type waters with high Li content can be derived
from the interaction between rocks and salt
solutions such as seawater and require no con
tribution from "magmatic" water. In this paper,lithium contents, and major chemical and stable
isotopic compositions of 30 saline waters in
Japan were determined to discuss the relation
ship between the lithium content and the genesis
of these saline waters.
SAMPLES AND EXPERIMENTAL PROCEDURES
Figure 1 shows the locality of stations for
water samples (Cl>5,000 mg/Q) used in this
study. Seawater for reference was collected off
Torishima Island, Tokyo (Takamatsu et al.,
1980). The temperature, alkalinity and pH
values of water samples were measured in situ.
Stable isotopes of oxygen and hydrogen in the
waters were analyzed on the mass spectrometer
in the Institute for Thermal Spring Research,
Okayama University. The isotopic ratios are
reported in 5180 and 51) values, relative to
SMOW.
Sodium, potassium and lithium were meas
2 Present address: Ocean Research Institute, University of Tokyo, Minamidai, Nakano-ku, Tokyo 164, Japan
143
144N. Takamatsu et al.
1 0
N Q
V:v
4
9
e 4"rDJAPAN
110SEA
268
23,25 20
CIP21
F~12
1711
.a
313
15
I
I
"Y10
i1822
-19
-1614
-27
29
24Fig.5
3 4
6
7 ~~---5
2
.0- 9
J4 PACIFIC OCEAN
O
Fig.. 1 Locations of the saline springs investigated.
Dotted area = Green-tuff region. F = Western margin of the Fossa Magna (Itoigawa-Shizuoka line).
M = Median Tectonic Line. Numbers accord with those of Table 1.
ured by atomic absorption spectrometry,
calcium and magnesium by EDTA titration
method or by atomic absorption spectrometry,
chloride and sulfate by ion chromatography.RESULTS AND DISCUSSION
Table I shows chemical and isotopic com
positions of saline waters investigated. It should
be noted that the lithium concentrations of the
Lithium in saline spring waters 145
Table 1. Chemical and isotopic compositions of saline spring waters investigated
No. Locality (Prefecture) pH Tw Li
(°C)Na K Ca Mg Cl SO4 HCO3 CU 6D 6180 Geology
(mg/e) (%(,)
Seawater (Tokyo) 0.15 10800 407 416 1320 19700 2700 1.0 0.0 0.0
1. Itoh (Shizuoka) 7.9
2. Kouzu (Tokyo) 7.6
3.Izusan (Shizuoka) 8.0
4. Ajiro (Shizuoka) . 7.8
S. Miyake (Tokyo) 6.9
6. Shikine (Tokyo) 5.5
7. Shimogamo (Shizuoka) 8.2
8. Katayamazu (Ishikawa) 7.3
9. Hachijo (Tokyo) 7.4
10. Matsunoyama (Niigata) 7.3
11. Moritake (Akita) 7.9
12. Tamugimata (Yamagata) 6.5
13. Yunosawa (Aomori) 7.1
14. Hikage (Akita) 6.5
15. Yatate (Akita) 6.3
16. Takizawa (Fukushima) 6.8
17. Yudonosan (Yamagata) 5.8
18. Isobe (Gunma) 7.6
19. Kitashiobara (Fukushima) 7.6
20. Wakura (Ishikawa) 8.0
21. Teradomari (Niigata) 8.0
22. Atsushio (Fukushima) 7.3
23. Arima-A
24. Ishibotoke
25. Arima-B
26. Takarazuka
27. Yashio
28. Aokura
29. Shionosawa
30. Kashio(Hyogo)
(Osaka)
(Hyogo)
(Hyogo)
(Gunma)
(Gunma)
(Gunma)
(Nagano)6.4
6.7
7.0
6.4
6.8
6.7
6.4
7.939.5 0.15 8470 146
39.0 0.39 9250 416
67.3 0.22 2090 32
78.8 0.33 3580 80
51.9 0.71 9750 430
64.5 1.07 8500 580
100 0.61 4080 280
72.5 1.11 3117 73
65.0 3.10 9400 810
74.7 1.72 3580 116
64.2 3.78 6030 226
23.8 4.32 4530 17
53.0 2.73 3400 201
42.2 2.40 2650 343
28.5 3.58 3860 382
55.5 4.81 6460 898
50.6 4.50 2980 557
24.5 8.22 10800 219
47.5 8.11 5960 208
87.0 8.26 4170 128
33.5 7.34 6610 41
59.0 4.64 2600 119
98.2 43.4 16800 3240
19.1 13.3 5580 69
21.5 28.2 9000 955
15.3 14.8 4970 532
19.8 19.2 7340 693
20.1 18.4 4870 385
14.5 15.5 3480 169
16.6 62.2 11540 1461220
397
3380
3773
816
661
2307
2150
1289
1870
2710
3394
518
627
745
692
1757
137
1250
3030
126
939
3020
960
1610
963
330
365
495
6611028 16340
1175 17710
98 8586
91 12290
995 17720
990 18820
7 10370
5 8860
409 18800
0.2 8828
0 13780
130 12530
213 6549
165 5554
429 7977
652 10700
250 8290
47 13070
276 12100
22 12050
90 9149
182 5730
10 34700
416 9260
19 17400
58 8900
86 10200
124 8237
155 6042
101 188002000 104
232 119
1053 31
648 20
152 328
1912 113
153 20
282 32
821 211
41.0 145
235 1340
0 226
9 107
11 758
2 1468
1250 2351
451 139
15 7560
678. 367
200 19
0 3004
354 127
0 36
1 3641
7 1270
0 2111
1090 3155
28 879
74 580
0 1041.2
2.9
3.4
3.5
5.3
7.5
7.7
17
22
26
36
45
55
56
59
59
71
83
88
90
110
110
160
190
210
220
250
290
340
440-17.8
-2.6
-42.8
-27.1
-9.6
-9.3
-22.4
+0.6
(-24.8)'
(-34.2)3
-54.1
(-44.2)3
(-39.7)3
-39.8
-62.2
-31.0
-64.2
(-17.6)2
-17.9
-67.9
(-37.9)5
(-44.8) 5
(-45.0)'
(-46.1)5
(-20.6)4
-51.4
-61.7
-53.3-1.4
+0.7
-5.6
-4.4
-0.6
+0.1
-2.5
+1.0
(+ 1.1)'
(-4.7)3
-8.3
(-3.8)'
(-2.0)3
-2.5
-10.2
+2.7
-8.6
(-4.7)2
-0.4
-10.7
(+4.4)1
(-3.9)5
(-2.4)5
(-5.3)5
(+6.0)4
-4.3
-7.5
-3.7Ba Q
Rh Q
An Q
Ba N
Ba Q
Rh Q
An N
Q
An Q
N
Vo T
An N
Vo T
Vo T
Vo T
Vo N
An Q
Q
N
N
N
N
Gr pC
Gr pC
Gr pC
Gr pC
Sc L
L
L
Sc L
-: not determined
Samples offered
Samples were offered by K. Aikawa (Nos. 2, 5, 6, 9, 10 and 21), by Y. Sakai (No. 18) and by H. Masuda (No.
24).
Isotopic analyses
1: after Abe and Sakai (1983), 2: after Matsubaya et al. (1973), 3: after Matsubaya et al. (1975), 4: after Sakai
and Matsubaya (1974), 5: determined by H. Masuda.
Abbreviations for the associated rocks (after Sumi, 1975)
pC: Precenozoic, L: Paleozoic, N: Neogene, Q: Quaternary, T.• Tertiary, An: Andesite and propylite, Ba: Basalt
and dolerite, Gr: Granite, granite porphry, quartz diorite, quartz porphyry, granodiorite, granophyre and quartz
syenite, Rh: Rhyolite and dacite, Sc: Schist, hornfels and phyllite, Vo: Volcanic rocks.
waters fluctuate widely and most of them are
much larger than those of seawater (0.15 ppm)
and river waters (0.1 400 ppb, Heier and
Billings, 1969) in the world.Problems for temperature dependency of lithium
contents of saline waters: Fouillac and Michard
(1981) have proposed a Na/Li geothermometer
for natural waters. They found two empirical
146 N. Takamatsu et alL
thermometric relationships between log (Na/Li)
and the inverse of absolute temperature. One is
for waters of low to moderate salinity (Cl<
0.2 M; they named it "general line") and
another one is for marine waters and brines
(Cl>0.3 M; they named it "brine line"). We
examined the relationship between log (Na/Li)
and temperature for our saline spring waters as
shown in Fig. 2. The underground temperatures
of most waters were estimated by the Na-K-Ca
method (Fournier and Truesdell, 1973) and
those of waters which seem to be contaminated
by seawater were calculated by the Mg correc
tion method (Fournier and Potter, 1979).
4
3
c z
0
2brine line
9 •1group 1
4
11 8
13 *15
11 0 • 18
1 15
group 4 /1028.62910 8
20.1
200 0
5 1
3
6• •,
group 2
i
12 C10
21
,.
A ~general line
g roup 31001
2 10
/T 3
Fig. 2. Relationship between log(Na/Li) values and
estimated temperatures of the saline waters. Open cir
cles are the waters under 35°C sampling temperature.
Numbers accord with those of Table 1.
Figure 2 shows no clear relation between
log (Na/Li) and the temperatures estimated .
The degree of deviation from the brine line of
the low temperature waters (generally contain
high HCO3) at the orifice seems to be larger than
that of higher temperature ones. Two major
reasons are considered as follows. One is in
accuracy of the geothermometers. Up to the
present, no geothermometer has been presentedwhich is universally valid for low temperature
and high Pco 2 waters, although an attempt was
made (Paces, 1975). The temperatures of
those waters may be estimated lower than the
actual underground ones, because they are in
disequilibrium with rocks. The other reason of
the deviation is that the waters had been en
riched in lithium in high temperature hydro
thermal systems and the lithium is still main
tained in solutions or trapped as leacheable salts
in sediments after cooling.
In any case, it can be concluded that the
lithium contents of saline spring waters in Japan
are not unequivocally controlled by a present
day temperature of the hydrothermal system,
but are associated with the genesis of the saline
springs as will be discussed below.
Li enrichment of saline spring waters: The
lithium concentrations of most saline waters are
larger than that of seawater as mentioned above.
Although it is uncertain whether all waters are
originated from seawater or not, a deviation
coefficient defined by CLi = (Li/Cl)sp/(Li/Cl)sw
(sp = spring water, sw = seawater) was calculated
to reveal the relative enrichment of lithium for
each saline spring water. The CU values are
listed in Table 1 and the relationship between
CLi and Cl content is shown in Fig. 3. The re
lationship between S D and S 18O of the saline
spring waters is shown in Fig. 4. The saline
springs can be classified into the following 4
groups by the relationships given above and the
regions from which the spring waters issue.
Group 1: Saline springs near coast. (Nos. 1 to 7
and No. 9 )
Group 2: Saline springs from Green-tuff region.
(No. 8 and Nos. 10 to 22)
Group 3: Saline springs from Osaka basin.
(Nos. 23 to 26)
Group 4: Saline springs from outer part of the
Median Tectonic Line. (Nos. 27 to 30)
Group 1 : The CLi values of the saline spring
waters occurring near coast were extremely low
(but larger than unity). The isotopic composi
tions of these waters are close to that of SMOW,
Lithium in saline spring waters 147
or are intermediate between seawater and local
meteoric waters. Thus these waters are con
sidered to have been formed by the interaction
of rocks with the waters which contain more or
less present-day seawater.
Takamatsu et al. (1983) found a clear cor
relation between CLi and CMg (similar deviation
coefficient for Mg) of the coastal thermal waters
in Izu islands and concluded that lithium is con
centrated by the ion exchange mechanism in the
thermal solutions with progress of rock-seawater
interaction. Similar discoveries have been made
in various submarine hydrothermal solutions
(Corliss et al., 1979; Von Damm et al., 1985a,
b). Von Damm et al. (1985a) reported on the
hydrothermal solutions from East Pacific Rise
at 21° N latitude that the lithium concentra
tions based on extrapolation of the data to zeroMg are about 6-9 ppm and these represent
enrichment factors (presumably identical with
CL) of up to 50 times the ambient values. These
findings evidently show that the lithium in these
waters are extracted from solid rocks in the
process of reactions of rocks with seawater or
diluted seawater. However, it is questionable
whether or not the lithium concentration of the
waters of group 1 is related to a chemical
equilibrium between water and a lithium mineral.
We support the idea that the linear relationship
between log (Na/Li) and the temperature
estimated for group 1 (cf. Fig. 2) reflects mainly
increasing degree of rock dissolution with in
creasing temperature, as suggested by Fouillac
and Michard (1981).
Group 2 : Many highly saline waters (5.55
500
400
300
J
U200
100group 40
Kashio
0 Shionosawa
group 20 Aokura
Yashio
0
Takarazuka
tshibotoke
A A group 3
Arima-B
Arima-A
group 1Seawater
Fig. 30 0.2 0.4 0.6 0.8
CI (ep,)
Relationship between CLi values and CI contents of the saline waters.
• : Group 1 A : Group 2
: Group 3 .0 : Group 41.0
148N. Takamatsu et al.
0
(0'20
6D=86180+10
-10 -510: 5 6180
.
2
r5
A
20
1.
1.
1A
21-20
100
27
4
18
11 1615
123 13
26 024
28 o..
25-4023
30
r
0--60
1929
22
Fig. 4 Relationship between liD and 5180 values of saline spring waters investigated. Symbols are the same
as in Fig. 3. Numbers accord with those of Table 1.
13.78 Cl/Q) issue out from Green-tuff regions,
especially from tectonically quiet sites during
the Quaternary (the Japan Sea side of Honshu).
However, some (e.g., No. 17: Yudonosan) occur
in proximity to active Quaternary volcanoes.
The isotopic compositions of many saline waters
of group 2 are similar to those of local meteoric
waters, but some waters deviate from the
meteoric water line (Fig. 4). The deviation
might be caused by the reaction of carbonate
minerals with meteoric waters at relatively high
temperature (Clayton et al., 1966; Hitchon and
Friedman, 1969). In any case, their high values
of salinity are considered to be produced by
dissolution of salts in the Green-tuff formations
which are composed of marine clastic sediments
of the Miocene age. The Na/Li ratios and CU
values of the waters of group 2 resemble one
another (cf. Figs. 2 and 3), regardless of their
temperature. This shows that they are notrelated to the temperatures of present-day hydro
thermal systems but may reflect the composi
tions of leacheable salts trapped in the sediments.
Relatively low CU values of some waters from
Green-tuff regions are probably due to the
contamination of saline brine by present-day
seawater because those occur near coast (e.g.,
No. 8: Katayamazu). The CLi values of the
waters of group 2 would fall within a relatively
narrow range of about 50 to 100, if we ignore
the data at No. 8 as the contaminated one. It
can be concluded that the lithium, as well as
chloride of these waters, may have originated
from the salts trapped in marine clastic sediments
and the lithium concentration of salts may
almost have been determined by the initial
genetic conditions of Green-tuff formations.
Group 3 : The definitions by Matsubaya and
Sakai (1973) of the formation mechanisms of
Lithium in saline spring waters 149
highly saline waters in Japan, as for Arima type
thermal spring waters, are not so clear. From
the relationship between CI content and 6 D or
5180 values, Matsubaya (1981) proposed that
only four brines in Japan belong to Arima type;
Arima, Takarazuka, Ishibotoke and Kashio
brines. Recently, Hashizume (1984) claimed
that Kashio brine is different from the other
three brines in chemical composition and that
the difference may be due to the difference of
host rocks of the saline waters. Therefore, this
paper regards the first three brines (Arima,
Takarazuka and Ishibotoke) as "the saline
springs from the Osaka basin." Figures 3 and 4
indicate that the waters from the Arima
Takarazuka-Ishibotoke area are formed by the
mixing of meteoric waters with the unique high
saline water having fairly high CU value (about
200). Although the formation mechanism of
the saline springs of this area has been debated
extensively (Miyake et al., 1954; Tsurumaki,1964; Matsubaya et al., 1973; Matsubaya et d,
1974; Tanaka et al., 1984; Masuda et al., 1985),
no decisive conclusion as to the origin of the
unique saline brine has yet to be drawn. The
fact that the CLi values of Arima and Takarazuka
(located on the northern rim of the Osaka basin)
are similar to that of Ishibotoke (on its southern
rim) is consistent with the hypothesis that the
brine waters of the Arima-Takarazuka-Ishibotoke
area were formed in marine sediments of the
Osaka basin and were later trapped into faults
and cracks developed by the continuous uplift
of the Rokko and Ikoma mountains and sub
sidence in the central area since the middle
Pliocene (Sakai and Matsubaya, 1977). Further
investigations on the genesis of the unique brine
of this area are needed to elucidate the lithium
enrichment for these waters.
Group 4 : Three
(Yashio, Aokurasaline brines of Gunma Pref.
and Shionosawa) and Kashio
F
G"3
'/ ,Ao1
kur
.
J /i. / r /i
/tii~I j/t%
/I I fill I
/K~shio'i $hi noswa'
J~OF OF
OF if OF," OF le OF OF
/ / / / ////if/
41 OF / OF Of/,Yashio
Mrte.:. i0 50 Km
Mt. FujiTokyo bay
Fig. 5 Geological map of saline springs from the outer part of Median Tectonic Line.
F = Western margin of the Fossa Magna (Itoigawa-Shizuoka line). M = Median Tectonic Line. S = San
bagawa metamorphic belt. P = Paleozoic formation. Me = Mesozoic formation. Ry = Ryoke metamorphic
belt. G = Green tuff formation.
150 N. Takamatsu et alL
brine of Nagano Pref. issue from the outer part
of the Median Tectonic Line. These brines are
characterized by low temperature (cf. Table 1)
and considerable oxygen isotopic shifts from
meteoric water line in Fig. 4. Yashio and Kashio
brines are in the Sanbagawa metamorphic crystal
line schist, while Aokura and Shionosawa are in
the Chichibu Palaeozoic formations adjacent to
the Mesozoic formations as shown in Fig. 5.
Tertiary sediments exist in the neighborhood of
these springs. However, the CLi values of these
waters are far larger than those of the waters
from the Green-tuff region. Sakai and Oki
(1978) have considered that the waters of
Yashio were confined in the fracture zone of the
crystalline schist when fossil seawaters in the
Green-tuff formation were extruded by the
crustal deformation. It seems reasonable that
the waters of Aokura and Shionosawa were also
formed by this mechanism. Hence, the high CLi
values of the three Gunma saline waters are
considered to be caused by the interaction be
tween the lithium bearing minerals in enclosing
rocks and the salt solution, if the origin of these
waters is the "fossil seawaters" which remained
in the Tertiary marine sediments.
The Kashio and Yashio brines are found in
the Sanbagawa crystalline schist near the Median
Tectonic Line (Hashizume, 1984) and hence a
similar mechanism for Li enrichment as Yashio
can be assumed. It is noteworthy that the CLi
value of Kashio of Nagano Pref. is the largest
among the saline waters of group 4, in spite of
the fact that Kashio is located far from the
Tertiary system as seen from Fig. 5. This implies
that the Kashio brine water is of older fossil
seawater which has strongly interacted with
rocks. containing high lithium. An accurate
measurement of the lithium content of rocks
related to the saline springs of group 4 will be
required to explain the high lithium content of
the waters of this group.
CONCLUSIONS
The Li concentrations of the saline waters in
Japan are not necessarily controlled by thepresent-day temperatures of the hydrothermal
systems, but are associated with the lithium con
tents of host rocks around the saline springs.
The CLi values of the waters are useful for the
genetic classification of the saline springs,
although detailed geochemical studies of lithium
and experimental works are still necessary. It is
worth noting that the inland saline waters
issuing out near the Tectonic Line have high CLi
values, in contrast with those issuing out near
the coastal line or with submarine hydrothermal
solutions from several sites along ridge crest
systems of the world.
Acknowledgements-The authors wish to express their
thanks to Professor K. Aikawa of Toho University and
Y. Sakai of Gunma Institute of Public Health for provid
ing water samples. Our sincere thanks are due to H .
Masuda of Osaka City University for her permission of
using her valuable isotopic data and for providing the
water sample. We also wish to thank Professor H. Sakai,
Institute for Thermal Spring Research, Okayama Univer
sity (presently Ocean Research Institute, University of
Tokyo), for his giving us an opportunity to use the mass
spectrometric equipments.
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Von Damm, K.L., Edmond, J.M., Grant, B., Measures,
C.I., Walden, B. and Weiss, R.F. (1985a) Chemistry
of submarine hydrothermal solutions at 21°N, East
Pacific Rise. Geochim. Cosmochim. Acta. 49,
2197-2220.
Von Damm, K.L., Edmond, J.M., Measures, C.I. and
Grant, B. (1985b) Chemistry of submarine hydro
thermal solutions at Guaymas Basin, Gulf of Califor
nia. Geochim. Cosmochim. Acta 49,2221-2237.
White, D.E. (1957) Thermal waters of volcanic origin.
Bull. Geol. Soc. Am. 68,1637-1658.
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Takamatsu (1986) saline spring waters of japan.txt
|
E-LETTER Earth Planets Space ,57, e21–e24, 2005
Gravity and density variations of the tilted Tottabetsu plutonic complex,
Hokkaido, northern Japan: implications for subsurface intrusive
structure and pluton development
Hiroyuki Kamiyama1, Akihiko Yamamoto2, Takeshi Hasegawa3, Takanori Kajiwara1, and Toru Mogi1
1Institute of Seismology and Volcanology, Graduate School of Science, Hokkaido University, Sapporo, Japan
2Department of Earth Sciences, Faculty of Science, Ehime University, Matsuyama, Japan
3Department of Earth and Planetary Sciences, Graduate School of Science, Hokkaido University, Sapporo, Japan
(Received September 9, 2005; Revised November 9, 2005; Accepted November 10, 2005)
An exposed cross section of the tilted Tottabetsu plutonic complex allows direct evaluation of its original
2-D cross-sectional shape and pretilting vertical density variations in both the pluton and the country rocks,which serves as a strong constraint in gravity modeling that complements information on the ‘missing’ pretiltinghorizontal dimension of this tilted pluton. The pluton is stratified with the uppermost thin granitic unit ( ∼1-
km thick) and the underlying thick gabbro-diorite units ( ∼9-km thick) that preserve a stratigraphic record of
numerous hotter replenishments in the form of alternation of originally horizontal mafic sheets and cumulatelayers. Both the pluton and the country rocks show systematic density increase with pretilting crustal depth, butdensity contrast of the pluton with the country rocks varies between each unit. The 2-D cross-sectional shapeand gravity analysis revealed that the pluton had a vertically-elongated shape with vertical side walls beforetilting. The vertical side walls, together with the stack of the originally horizontal sheets and cumulate layers,suggests that the pluton grew only vertically by piston mechanism. The very thick, exposed cross section providesunequivocal evidence for development of such a pluton with this unusual shape and mass distribution, which hasbeen inferred elsewhere only by some geophysical studies.
Key words: tilted pluton, pluton emplacement, gravity anomaly, rock density.
1. Introduction
Construction of a pluton is a multi-process geological
event including emplacement, growth and solidification ofmagma bodies within the Earth’s crust. Important informa-tion on the mechanism by which plutons are constructed ispreserved in their shapes and internal structures. However,exposures of the overwhelming majority of plutons onlyprovide sections that can be approximated horizontally, soinformation on the vertical variations in shape, lithologyand internal structures is severely limited.
Considerable interest has focused on geophysical data
bearing on subsurface structures of plutons (Ameglio andVigneresse, 1999). Especially, gravity techniques havewidely been applied to studies of plutons (e.g. Ameglio andVigneresse, 1999 and references therein). Although gravity(and other geophysical) data can place important constraintson overall subsurface geometry, their limited spatial reso-lution does not allow evaluation of outcrop-scale featuresof unexposed parts of plutons. Furthermore, lack of exactknowledge of vertical variation in density contrast betweenplutons and country rocks leads to unreliable gravity esti-mates of shape and depth extent of plutons (Pitcher, 1979).For many plutons and their country rocks, density variationshould be much more pronounced in the vertical than in the
Copy right c/circlecopyrtThe Society of Geomagnetism and Earth, Planetary and Space Sci-
ences (SGEPSS); The Seismological Society of Japan; The V olcanological Societyof Japan; The Geodetic Society of Japan; The Japanese Society for Planetary Sci-ences; TERRAPUB.horizontal dimension, but vertical variation in rock density
is seldom taken into account.
Another approach of geological study only can be pos-
sible in the field of rare vertical cross sections of plutonsexposed in tilted crustal sections. This allows direct ob-servation of cross-sectional shape and vertical variations inlithology and small-scale structures (e.g. Miller and Miller,2002). Nevertheless, it is still largely limited to 2-D surfaceexposures. Geophysical study would complement informa-tion on the ‘missing’ original horizontal dimension of suchtilted plutons. To our knowledge, however, none of the geo-physical studies have been applied to the tilted plutons.
An exposed cross section of the tilted Miocene Tottabetsu
plutonic complex (TPC) provides not only its 2-D cross-sectional shape, but also a stratigraphic record of crystal-lization and recurrent replenishments. Furthermore, thecross-sectional exposure allows direct evaluation of pretilt-ing vertical variation in density contrast between the plutonand the country rocks, which serves as a strong constraint ingravity analysis. Here we present field and gravity evidencethat describes the geometry, mass distribution and internalstructure of this tilted pluton, which suggest that the plu-ton grew only vertically by piston mechanism to form anunusually thick cylindrical pile of mafic rocks and an over-lying thin granitic cap.
e21
e22 H. KAMIYAMA et al. : GRA VITY AND DENSITY V ARIATIONS OF A TILTED PLUTON
Fig. 1. Geologic map of the Tottabetsu plutonic complex modi fied from
Suetake (1997) (a), and simpli fied maps of the Japanese Islands (b) and
south-central Hokkaido (c). Paleohorizontals are inferred from ma fic
sheets, pipes and cumulate layering and foliations.
2. Geologic Background
The Hidaka metamorphic belt (HMB), north Japan
(Fig. 1(b) and (c)), represents a tilted cross section throughpart of continental crust (Komatsu et al. , 1983). It is ex-
posed as a consequence of collision of Kuril fore arc andNortheast Japan arc due to southwestward migration ofthe former since the late Miocene (Kimura, 1996). Folia-tion and banding of the metamorphic rocks and layering ofthe plutonic rocks in this crustal section in general strikeroughly N-S and dip steeply eastward. Metamorphic gradesystematically increases westward from zeolite to granulitefacies, indicating that the crustal depth increases westward.The maximum pressure and temperature conditions were 7kbar and 830
◦C (Osanai et al. , 1991). A major east-dipping
thrust, the Hidaka main thrust (HMT), bounds the westernedge of the HMB (Fig. 1(c)).
The TPC (Fig. 1(a)), which constitutes part of the HMB,
intrudes metamorphosed equivalents of sedimentary rocksof the Nakanogawa group along its northeastern, easternand southern margins, and is surrounded by a contact au-reole that overprints greenschist- to amphibolite-facies as-semblages (Takahashi, 1992). The TPC is bounded on thenorthwest by the Pipairo tonalite, and on the west by theLower gabbro-diorite (LGD) that consists of gabbros andsubordinate diorites. The regional subsolidus deformationin the Pipairo tonalite and the LGD implies that they pre-dated the TPC, which is undeformed except for minor frac-turing.3. Cross-Sectional View of the TPC
The TPC is roughly rectangular in plan, measuring 8 km
by 12 km (Fig. 1(a)). There is little topographic control onthe map pattern of the complex, implying that the complexpresently has steeply dipping sides. Paleohorizontals in-ferred from magmatic structures (see below) roughly strikeN-S and dip steeply eastward in the range between 50
◦and
90◦(Fig. 1(a)). On the basis of the dips of the paleohor-
izontals and E-W extent of the complex, thickness of theentire complex is estimated to be ∼10 km. Width of the
complex is nearly constant, ∼8 km in exposed lateral (N-S)
extent, throughout the entire stratigraphic succession. Thepaleohorizontals are roughly perpendicular to the exposedoriginal side walls and subparallel to the original roof of thepluton. Thus, in terms of the 2-D cross-sectional exposure,the TPC apparently had a cylindrical form with vertical sidewalls and a flat roof before tilting.
Suetake (1997) divided the TPC, from west to east, into
zones I to III on the basis of dominant lithology of eachmappable area (Fig. 1(a)). Zone I ( ∼7-km thick) is pre-
dominantly composed of gabbroic rocks and occupies thelower two thirds of the complex. It is overlain by zone II(1–1.5-km thick), which consists of wide variety of dioritic
rocks. Zone III (1 –1.5-km thick) forms the granitic cap at
the uppermost part of the complex.
In zone I, steeply east-dipping sheets of fine-grained
gabbros are interlayered with and variably chilled againstmedium-grained gabbroic and leucodioritic cumulate lay-ers. The alternation of sheets and cumulate layers are takenas evidence for ponding of recurrently injected ma fic mag-
mas on the progressively aggrading crystal-rich base of amore evolved, crystal-poor magma chamber (e.g. Wiebeand Collins, 1998). Lobate basal margins of the gab-broic sheets, flame structures, and pegmatitic pipes pene-
trating the gabbroic sheets serve as indicators of ‘way-up ’
and horizontals at the time of their formation (Wiebe andCollins, 1998). The pipes have axes roughly perpendicularto the sheets and cumulate layers, indicating that the sheetsand layers were originally horizontal. Similar east-dippingsheets and cumulate layers are also seen in zone II, but therocks are distinctly more felsic, intermediate in composi-tion. The overall compositional difference between zonesI and II is essentially attributed to compositional differenceof the injected magmas. By contrast, zone III consists ofmassive homogeneous granite that presumably formed ascomplementary fractionates to the cumulates in zones I andII.
4. Density Variations
Densities of surface rock samples were determined by
laboratory measurements. Since the HMB represents thetilted cross section through lower to upper continentalcrust whose paleodepth increases westward (Komatsu et al. ,
1983), systematic E-W variations in rock densities might beexpected. For this reason, results of the density measure-ments were plotted versus E-W distance from the HMT thatdefines the deepest part of the crustal section (Fig. 2).
Signi ficantly, densities of the felsic country rocks includ-
ing metasedimentary and tonalitic rocks systematically de-crease eastward from ∼2.8 to ∼2.6 g/cm
3for over the dis-
H. KAMIYAMA et al. : GRA VITY AND DENSITY V ARIATIONS OF A TILTED PLUTON e23
Fig. 2. Rock densities plotted versus E-W distance from the Hidaka Main
Thrust.
tance of ∼30 km. As there is little correlation between
the measured densities and whole-rock chemical composi-tions, the eastward decrease in density of these felsic coun-try rocks appears to re flect change in proportion of con-
stituent minerals primarily due to the decreasing metamor-phic grade. Similar systematic eastward decrease in rockdensity is also observed for the TPC and the LGD. Densi-ties of the gabbros from zone I and the LGD gently decreaseeastward with concomitant increase in modal hornblende atthe expense of anhydrous ma fic minerals. However, density
contrast of these gabbros relative to the felsic country rocksat a given distance from the HMT is found to be nearly con-stant, ∼0.2 g/cm
3. The diorites in zone II show marked
eastward decrease in their densities, re flecting eastward in-
crease in SiO 2content. The average density contrast of the
diorites relative to the metasedimentary country rocks ofcomparable distance from the HMT is close to 0.1 g/cm
3.
Granites of zone III show negative density contrast with themetasedimentary country rocks. The average density con-trast of the granite relative to the metasedimentary countryrocks is close to −0.1 g/cm
3.
5. Gravity Anomaly and Subsurface Contact
Previous gravimetric studies (e.g., Yamamoto et al. ,
2001) do not illustrate detailed features of the TPC dueto lack of gravity data and nonuniformity of station cover-age. We performed gravity measurements such that stationcoverage would be as uniform and dense as possible in thestudy area. Finally, 250 gravity data were newly obtainedusing the Scintrex gravimeter with a precision of ±0.05
mGal. Elevations were determined using differential GPS,spot heights and contours on 1/25,000 topographic mapspublished by Geographical Survey Institute (Japan). In to-tal, 887 gravity data were reduced for free-air and Bouguercorrections as well as for terrain effects, using a density of2.67 g/cm
3. Terrain corrections were based on the method
of Yamamoto (2002).
Despite the limited distribution of gravity stations, the
Bouguer anomaly map does show some distinct correlationsbetween the exposed area of the TPC and local gravity field
impinged upon eastward declining regional gravity field
(Fig. 3). As expected from the positive density contrasts ofthe gabbros and diorites relative to the felsic country rocks,an east-plunging ridge of gravity high is seen associatedwith the exposure of zones I and II. At the same time, a
Fig. 3. Bouguer anomaly map of the Tottabetsu plutonic complex. Con-
tours are in milligals. The black dots mark the locations of gravity sta-tions. The pro file line (A-A
/prime) shown in Fig. 4 is indicated.
Fig. 4. Gravity interpretation along E-W pro file A-A/prime(Fig. 3). (a) Bouguer
gravity pro file and subtracted regional gravity. (b) Residual anomaly
and modeled gravity field. (c) 2-D model for the gravity pro file. Density
contrasts relative to the metasediments are shown in g/cm3.
weak gravity low is seen associated with the exposure of
zone III, concordant with the negative density contrasts ofthe granite relative to the metasedimentary rocks.
The Bouguer anomalies include the regional effects in-
duced by far-located or deep-seated sources unrelated tothe TPC. Speci fically, in the HMB, it must also include
the effect of the lateral variation in rock densities (Fig. 2).To remove these regional effects, long-wavelength anomalycalculated by running average of a square grid of 20 by 20
e24 H. KAMIYAMA et al. : GRA VITY AND DENSITY V ARIATIONS OF A TILTED PLUTON
km2was subtracted from the Bouguer anomaly (Fig. 4(a)).
The speci fic grid size of 20 by 20 km2was chosen in or-
der for the resultant residual gravity at the exposed area ofthe country rocks to be flat. Figure 4(b) shows the resid-
ual anomaly and the gravity field calculated from the 2-
D model shown in Fig. 4(c). The gravity feature of themodel closely matches the observed residual anomaly. Themodel shows that the subsurface contact between the TPCand metasedimentary rocks linearly deepens westward witha dip of 10
◦–20◦. Such a gently west-dipping contact is
roughly perpendicular to the steeply east-dipping paleohor-izontals inferred from the surface geology (Fig. 1), stronglysupporting that the hidden pretilting side wall of the TPCwas also nearly vertical.
6. Discussion and Conclusions
Ourfield observations suggest that the TPC was formed
by simultaneous crystallization and recurrent replenish-ments during the evolution of a single magma chamber.This in turn means that the active magma chamber at anyone time was much thinner than the entire complex.
When the magma chamber was replenished by new
magma, the magma chamber must have in flated in order to
accommodate the newly injected magma, space for whichmight be created by displacing country rocks aside, upward,or downward. The construction by gradual vertical stackingof the horizontal sheets does not match lateral displacementof country rocks as the mechanism of space creation. Thecountry rocks therefore must have been displaced either up-ward by roof lifting or downward by floor depression as
envisaged for some tabular intrusions (e.g. Cruden, 1998).We suggest that the space for the successive batches of hot-ter magmas was made by vertical displacement of fault-bounded blocks of roof or floor country rocks with little
horizontal displacement. Evidence consistent with this pis-ton mechanism includes the originally vertical side wallsandflat roof and near-constant attitudes of the paleohori-
zontal structures that were formed during deposition on thechamber floor.
According to recent compilations (e.g. Petford et al. ,
2000), most plutons are tabular with vertical dimensionmuch shorter than horizontal ones, though a few works ontilted plutons has documented more equant shapes (Haeus-sler and Paterson, 1993; Bachl et al. , 2001). Furthermore,
strong positive correlation between horizontal and verticaldimensions in the compilations implies that plutons grow byboth lateral spreading and vertical in flation as a general rule.
However, the shape and growth style of the TPC clearly donot match this empirical rule. The TPC is unusual in thatit has a cylindrical shape with vertical dimension slightlygreater than horizontal one and was shown to have grownonly by vertical in flation.
The TPC represents, to our knowledge, the only docu-
mented cross-sectional exposure of such a pluton with anunusually thick, cylindrical mass of ma fic rocks overlain by
a thin granitic cap. Plutonic complexes with similar shapeand mass distribution to the TPC have also been inferredelsewhere, but only by geophysical studies (Bott and Tu-son, 1973; Bott and Tantrigoda, 1987; Bauer et al. , 2003).
The very thick, exposed cross section of the TPC providesunequivocal evidence for development of such a pluton with
the unusual shape and mass distribution.
Acknowledgments. We thank Hitoshi Mori and Hiroaki Taka-
hashi for assistance in GPS measurements, Takashi Nakajima formany discussions, and Tadashi Usuki for providing part of rocksamples. We would also like to thank Calvin Miller for helpfulcomments on the earlier version of the manuscript, and MasaoKomazawa and Roman Teisseyre for reviews and comments thathelped us clarify the manuscript.
References
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Bachl, C. A., C. F. Miller, J. S. Miller, and J. E. Faulds, Construction of
a pluton: evidence from an exposed cross section of the Searchlightpluton, Eldorado Mountains, Nevada, Geol. Soc. Am. Bull. ,113, 1213 –
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crustal model of intrusive structures beneath the Messum ring complex,Namibia, Earth Planet. Sci. Lett. ,216,6 5–80, 2003.
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netic anomalies over the Mull Tertiary intrusive complex, NW Scotland,J. Geol. Soc. Lond. ,144,1 7–28, 1987.
Bott, M. H. P. and J. Tuson, Deep structure beneath the Tertiary volcanic
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blocks, Eldorado Mountains, Colorado River Rift, NV , USA, Lithos ,61,
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Geol. Soc. Lond. ,136, 627 –662, 1979.
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J. Phys. Earth, 26, Suppl., S 367-S 378, 1978
MAGNETIC STRATIGRAPHY OF THE JAPANESE NEOGENE
AND THE DEVELOPMENT OF THE ISLAND
ARCS OF JAPAN
Nobuaki NIITSUMA
Institute of Geoscience, Shizuoka University, Shizuoka, Japan
(Received June 19, 1978; Revised September 11, 1978)
Detailed correlation and chronology of the Neogene and Quaternary marine sediments
in the Japanese island arc area have been established by the application of magnetostrati-
graphic methods supplemented by microbiostratigraphic data. The sedimentation rates of
several sedimentary sequences show a similar pattern among the studied areas, and two
drastic synchronous changes in the rates of sedimentation are recognized. Thus the
Japanese Neogene and Quaternary can be divided into three major time intervals, named
I, II, and III in increasing age. The boundaries between these three intervals are 4.7
mybp (base of the Gilbert Epoch; magnetic anomaly3) and 10.4mybp (Epoch 9; magnetic
anomaly 5). The geographic distribution of the land area during the time interval I and
II was similar to the present; however, in the time interval III, it is completely different
but similar to the present Bonin-Mariana arc area. It has been documented by Hawaiian
hot spots and spreading features on the East Pacific Rise that the plate motion in the Pacific
Ocean area has also changed drastically. The time interval I is the period of high rate of
sedimentation (several hundreds cm/1,000 years) and moderately increasing plate motion;
the time interval II extremely low rate of sedimentation (less than several cm/1,000 years)
and slow plate motion, and at the same time land areas were expanded; the time interval
III moderate rate of sedimentation (several tens cm/1,000 years) and high rate of plate
motion, and land areas were reduced. These drastic changes can be explained by the"cyclic evolutionary model", originally proposed by Kanamori, and Forsyth and Uyeda's
slab-pulling driving force of the oceanic plate motion as follows. The drastic change from
the time interval III to II is ascribable to detachment of the down-going slab from the ocean
plate. The reduction in plate motion may also be triggered by the detachment, which re-
leases the ocean plate from the down pulling force. The high rate of sedimentation in the
time interval I is resulted from the steepening in the topographic relief and the increase in
the amplitude of tectonic deformation, which should be related with the horizontal com-
pressional stress caused by the coupling between the continental and oceanic lithospheres.
1. Introduction
The island arc systems, now found mainly along the continental margins around the
Pacific, are particularly important areas to understand the geodynamics of the world.
It has been conclusively documented that the oceanic lithosphere bends and descends
into the mantle at the island arc system. In the theory of plate tectonics, lithospheric
subduction is one of the major tectonic processes related to the formation and evolution
of island arcs. There is a diversity of topography, seismicity, Wadati-Benioff zones,
volcanism, gravity anomalies, stress fields, etc., in the island arc systems. It is almost
certain that the degree of mechanical coupling between the oceanic and continental
lithospheres varies among different island arcs, and the coupling may play an essential role
in the diversity of the systems (KANAMORI, 1977).
S 367
S 368 N. NIITSUMA
There are two kinds of model which have been proposed to explain the diversity and
evolution of island arc systems:
1) Stationary subduction model-inter-plate coupling is controlled by relative
movement of both plates, and the relative movement is quasi-stationary or does not have
any special trend of change;
2) Non-stationary subduction model-inter-plate coupling changes with evo-
lutionary stages of plate subduction (the evolutionary model of KANAMORI, 1971, 1977;
and KOBAYASHI and ISEZAKI, 1976).
If some changes in the tectonic features of an island arc system during geologic time
are found, they may provide us with some constraints on the island arc models. For such
an investigation, the Japanese island arcs area is most suitable because of the existence of
ample geologic and geophysical data. Unraveling the geological records and determining
the trends in tectonic features requires knowledge of the precise chronology of relevant
geological events, which can be compared with the history of the oceanic plate motion.
For this purpose, the magnetic stratigraphy is the most useful method because:
1) magnetic stratigraphy can be used on a gloval scale;
2) magnetostratigraphic patterns can be directly correlated with ocean magnetic
anomaly patterns, which show the spreading history of an ocean plate.
In this article, the magnetic stratigraphy of Japanese Neogene and Quaternary sedi-
ments and the tectonic evolution of the Japanese island arcs during the last 15 million years
are discussed, referring to the mechanism of Pacific plate motion.
2. Magnetic Stratigraphy of Japanese Neogene and Quaternary Sequences
Japanese Neogene and Quaternary rocks have been investigated paleomagnetically
since MATUYAMA (1929) recognized periodic change in the polarity of the geomagnetic
field. KAWAI (1951) applied paleomagnetic methods to sedimentary rocks, and reported
several horizons in which remanent magnetization. had reversed polarity. NAKAGAWA
et al. (1969) first applied the combined magneto-and biostratigraphic techniques to sedi-
mentary sequences, and found that a correlation of the sedimentary sections with the deep-
sea bottom sediments was possible. As the next step, magneto-and biostratigraphic in-
Fig. 1. Index map of investigated areas in the
Japanese island arc system.
Magnetic Stratigraphy of the Japanese Neogene and the Development of the Island Arcs S 369
vestigations on Japanese Neogene and Quaternary sediments outcropping in the Akita,
Sendai, Niigata, Choshi, Boso, Shizuoka, Miyazaki, Kumejima, and Miyakojima areas
have been made by the author and his colleagues (NIITSUMA, 1970, 1976; KIMURA, 1974;
NAKAGAWA and NIITSUMA, 1977; NAKAGAWA et al., 1975, 1977; NITOBE, 1977). Figure 1
shows the examined areas, which are widely distributed in the Northeastern Japanese arc,
Southwestern Japan, and the Ryukyu arc.
Figure 2 is an example of magneto-and biostratigraphy in the Boso area, central
Japan. The vertical axis corresponds to the thickness of the sedimentary sequence, and
the columns show lithostratigraphic classification (Formation), geologic columnar section
marked with the horizons of key tuff layers, intensity of remanent magnetization after
cleaning with alternating field and thermal demagnetization methods, latitude of virtual
geomagnetic pole position (VGP) calculated from the direction of remanent magneti-
zation, geomagnetic polarity determined by the latitude of VGP, and magnetostratigraphic
classification (Magnetozone). On the right-hand side of this figure, the alphabetically
nominated horizons correspond to the microbiostratigraphic data, which are shown in the
figure caption. The Boso area has a sequence of several thousands meters thick continuous
marine sediments, which is probably one of the thickest Neogene successions in the world
of which geology, paleontology, sedimentology, and tephra-stratigraphy have been well
studied.
The magnetostratigraphic sequence can be correlated with the standard normal and
reversed geomagnetic polarity pattern, supplemented with the ocean magnetic anomalies
(Fig. 3), using the microbiostratigraphic datums which are shown as the alphabetic nomi-
nations in Fig. 2 and beside the left column in Fig. 3. Those microbiostratigraphic datums
have been calibrated to the standard magnetostratigraphic scale (RYAN et al., 1974; SAITO
et al., 1975; TAKAYANAGI et al., 1978), as shown the alphabetic nominations beside the right
column in Fig. 3. This correlation makes it possible to calibrate the sedimentary sequence
to the absolute age. The 5,500m thick sedimentary sequence in the Boso area is defined
to be deposited for the last 16 my, and a one-third of the duration corresponds to the
remarkably thin part of the sequence (80m in thickness; magnetozones of BS-H and
BS-I) as shown by the correlation line in Fig. 3. This fact shows the existence of drastic
changes in the rate of sedimentation in this area, and in this part of the sequence, the rate
of sedimentation is extremely low. Using the calibration, the rate of sedimentation can be
calculated as shown in Fig. 5. In this part (4.7 to 10.4mybp), the sedimentation rate is
ranged from 1.0 to 10cm/1,000 years and increasing in the rate appears in the upper hori-
zons of this part. Above this part, the sedimentation rate is ranged from several tens to
several hundreds cm/1,000 years. Below this part, the rate is almost constant, ranged from
20 to 40cm/1,000 years.
Based on the geologic sequence in the Boso area and these changes in sedimentation
rates, the author proposes the following geochronologic classification of tectonic evolution:
i.e. the time intervals I, II and III (increasing in age). The boundaries between these
three time intervals are 4.7mybp (basal part of the Gilbert Epoch; magnetic anomaly 3)
and 10.4mybp (Epoch 9; magnetic anomaly 5).
The same kind of magnetostratigraphy in other areas has also been established. Fig-
ure 4 shows magnetostratigraphic correlation with supplemented microbiostratigraphic
data. Based on this correlation, the rate of sedimentation of those sections can also be
calculated as shown in Fig. 5. It can be seen that the previously mentioned drastic change
in sedimentation rates occurred not only in the Boso area but also in the other areas.
S 370 N. NIITSUMA
Fig. 2. Magnetostratigraphy of Neogene and Pleistocene sedimentary sequence in the Boso area
and selected microfossil events (NIITSUMA, 1976; NAKAGAWA and NIITSUMA, 1977; NAKA-
GAWA et al., 1977; TAKAYAMA, 1973, 1976; ODA, 1975. A, Lowest occurrence of Gephy-
rocapsa oceanica; B, highest occurrence of Discoaster spp. and coiling change of Pulleniatina
from left to right; C, lowest occurrence of Globorotalia truncatulinoides; D, coiling change of
Pulleniatina from right to left; E, highest occurrence of Sphaeroidinella spp.; F, coiling change
of Pulleniatina from left to right; G, highest occurrence of Globorotalia margaritae; H, highest
occurrence of Globigerina nepenthes; I, lowest occurrence of Pulleniatina primalis; J, lowest
occurrence of Globorotalia acostaensis; K, highest occurrence of Discoaster cf. hamatus; L, lowest
occurrence of Globigerina nepenthes; M, lowest occurrence of Orbulina suturalis.
Magnetic Stratigraphy of the Japanese Neogene and the Development of the Island Arcs S 371
Fig. 3. Magnetostratigraphic correlation with the
standard polarity pattern, supplemented with
the ocean magnetic anomalies (RYAN et al.,
1974), and geochronologic classification of tec-
tonic evolution in the Boso area.
Additional to these results, similar changes in sedimentation rate are found in DSDP site
438 (leg 57) taken from the inner trench slope of the northern part of the Japan trench
system (VON HUENE et al., 1978). It can be suspected that these drastic changes in sedi-
mentation rates have some implications in the tectonic evolution of the Japanese island arcs.
3. Tectonic Evolution of the Japanese Island Arcs in the Last 15 my
In this section, the tectonic evolution of the Japanese island arcs in the last 15 my is
discussed, based on the precise chronology and geochronologic classification by magneto-
stratigraphy mentioned above. The geologic character of the sedimentary sequences
in the Japanese island arcs is mentioned below for each time interval:
Time interval I (0 to 4.7 my). The sedimentation rate is high (several hundreds cm/
1,000 years) in the coastal plain area and intramontane basins. Distribution of marine
sediments is limited to the coastal area.
Time interval II (4.7 to 10.4my). In the coastal areas such as the Boso, Akita, Niigata
and Shizuoka areas, the sedimentation of marine sediments occurred and the rate of sedi-
mentation is extremely low (less than several cm/1,000 years). In the inland area, a large
amount of terrestrial pyroclastic deposits is supplied by dacitic volcanism, and the dis-
tribution of this volcanism is limited to the zone of the present Backbone Ranges (KITA-
MURA and ONUKI, 1973).
Time interval III (10.4 to 16my). This time interval is characterized by intermediate
sedimentation rates, the widest distribution of marine sediments (KITAMURA and ONUKI,
1973), and the Kuroko deposits which are believed to be created by submarine volcanic
activity (HORIKOSHI, 1976).
Figure 6 shows the paleogeographic maps for these time intervals. Since the early
stage of time interval I, the distribution of land area is similar to the present (Fig. 6 (a)).
In the time interval II (Fig. 6 (b)), also, it has a similar pattern except in the Ryukyu arc
(NAKAGAWA et al., 1977). In time interval III, it has a quite different pattern from the
S 372 N, NIITSUMA
Magnetic Stratigraphy of the Japanese Neogene and the Development of the Island Arcs S 373
Fig. 5. Changes in rate of sedimentation of Neogene and
Pleistocene sedimentary sequence and geochronologic
classification of tectonic evolution in the Japanese island
arc area.
present (Fig. 6 (c)). The volcanic activity is marked in Fig. 6 by crosses (~). In time
interval II, volcanism is characterized by terrestrial dacitic activity. In the later part
of time interval III, volcanism is characterized by basalt and dolerite extrusion (KONDA,
1974). The extrusion is assumed to have occurred under tensional conditions (KOBA-
YASHI and NAKAMURA, 1978). The paleogeographic features in time interval III are
rather similar to those of the Bonin-Mariana island arc system where the island arc con-
sists mainly of non-volcanic islands such as Ogasawara and the Saipan islands, and the
submarine volcanic activity occurring.
4. The Condition of the Lithospheric Plate under the Northeastern Japanese Island Arc and the
Cyclic Evolutionary Model
In the present Northeastern Japanese island arc system, the continental and the
oceanic lithospheres are nearly completely decoupled so that no major thrust earthquake
can occur along the interface, according to KANAMORI (1977). Because of the reduced
coupling, the tensional force caused by the gravitational pull of the denser downgoing
lithosphere may be transmitted to the oceanic lithosphere and may cause a large intra-
plate normal fault earthquake. The 1933 Sanriku Earthquake is interpreted as a litho-
spheric normal-fault which cuts through the entire thickness of the lithosphere (KANAMORI,
1977).
On the basis of the seismological results, UYEDA and KANAMORI (1978) estimated that
the subduction type in the Northeastern Japanese island arc system is an intermediate one
between the two end members. The end members are named the Chile and Mariana
types. In the Chile type, the coupling between the lithosphere of ocean and continent is
strongest and the width of the contact zone is very large, so that when slippage occurs it
results in a really major earthquake. This type is also characterized by the shallow angle
S 374 N. NIITSUMA
Fig. 6. Paleogeographic maps of the Japanese island
arc area in the each tectonic evolutionary time
interval.
of the Wadati-Benioff zone and the lack of deep earthquakes below 200km. In the Mari-
ana type, the coupling is weakest, and decoupling and detaching may occur. This type
is characterized by the steep angle of the Wadati-Benioff zone and by an active back arc
basin.
KANAMORI (1971, 1977) presented an interesting explanation for these features in
terms of gradual decoupling at the interface zone between the landward and subducting
plates. The Mariana type corresponds to the last stage of the evolutionary trend.
KOBAYASHI and ISEZAKI (1976) modified the original Kanamori's model to be more
consistent with geophysical observations such as magnetic anomalies, thickness of sediment
cover and age of the marginal sea floor. The modification is based upon the assumption
that subduction always starts at the boundary between the continental and oceanic litho-
spheres. (In the case of the original model, the oceanic plate, having lost the mechanical
support of the opposing continental lithosphere, may start sinking from the leading edge.)
When this process is completed, the detached slab of oceanic lithosphere sinks down to the
deeper asthenosphere, and the new cycle will repeat by renewed subduction.
The observations of the spatial distribution of earthquake hypocenters in the North-
eastern Japanese arc system suggest that the Wadati-Benioff zone reaches 550km in depth
Magnetic Stratigraphy of the Japanese Neogene and the Development of the Island Arcs S 375
and has 30 inclination (ISHIDA, 1970). Using these values, the length of the Wadati-
Benioff zone can be calculated as 1100km. If we use the average moving rate of the Pacific
plate of 10cm/year, estimated by Hawaiian hot spot data (JACKSON, 1976), it can be cal-
culated that 1,000km plate slab needs only 10my for sinking. This could be interpreted
to mean that the present lithospheric plate under the Northeastern Japanese island arc
system started to sink about 10mybp. From this, we can examine the two kinds of sub-
duction models, stationary and non-stationary, using the relationship between subducting
slab and tectonic evolution of the island arc system in the last 15my.
If we use the modified Kanamori's model, called the "cyclic evolutionary model"
in this article (Fig. 7), we should find the geologic events corresponding to the beginning
of the present cycle and previous cycle in the last 15my, based on the chronologic data.
Using the above-mentioned geologic and paleogeographic evidence, we can, say that the
present status goes back to the time interval II. It seems that the present cycle of plate
subduction started in the early part of the time interval II, 10mybp. This estimated age
agrees well with the possible starting time of sinking of the present plate under the island arc
system. Before the beginning of this cycle in the time interval II, the island arc system
should be in the last stage of the previous cycle, similar to the Mariana type. It has
already been noticed that the paleogeographic map of the time interval III shows a similar
pattern of land area distribution to the Bonin-Mariana arc. Because of the consistency
with the chronologic and paleogeographic data, we can conclude that the "cyclic evolu-
tionary model" is able to explain the process of oceanic lithospheric subduction.
Figure 7 illustrates the simplified cycle of the "cyclic evolutionary model," referring
to the development of the Northeastern Japanese island arc system. In this figure, (e)
shows the present status of the Northeastern Japanese island arc system. In (a) (early
time interval III), we appear to have a configuration similar to a slightly later develop-
ment of the situation shown in (e). The subductive slab increases in length and the gravi-
tational down-going force increases. This affects the decreasing land area in the island
arc system. Because the bending portion is pulled toward the oceanic side, the back arc
area is under a tensional stress field, by the suction force between continental and sub-
ductive slab (ELSASSER, 1971). The inclination of the slab increases to 45 or 50. This
stage corresponds to the present Izu arc. In (b), the slab becomes heavier and its incli-
nation increases up to 90. The main part of island arc subsides below sea level. The
bent portion of the slab moves toward the ocean, and because the relative position of the
trench and island arc is kept more or less constant by the suction force, the back arc basin
opening, which occurred under tensional stress field, is associated with extrusion of igneous
rocks and Kuroko deposits. This stage corresponds to the present Mariana arc, and the
later stage of the time interval III of the Northeastern Japanese arc. In (c), the down-
going slab is detached from the oceanic lithosphere. The island arc system is released
from the heavy lithospheric slab, isostatic rebound should occur and also release from the
horizontal stress. A new cycle is initiated and the oceanic lithosphere begins to subduct
between continental and oceanic lithospheres. This stage corresponds to the early part
of the time interval II. In this stage (d), the oceanic lithosphere underthrusts beneath the
continental lithosphere and is opposed by the latter. Because of its strength, the oceanic
lithosphere is unlikely to bend and low-angle (10 to 20) thrusting occurs. The stress
in the oceanic and continental lithospheres is compressive. At this stage, caused by the
strong coupling between the continent and oceanic lithospheres, major earthquakes occur.
This stage corresponds to the present Chile area and to the later stage of the time interval
S 376 N. NIITSUMA
Fig. 7. Schematic "cyclic evolutionary model" for the
plate subduction in Northeastern Japanese island arc
system and geochronologic classification of tectonic
evolution.
II and early part of the time interval I.
5. Discussion and Conclusions
The relationship between the sedimentation rate in the Japanese sedimentary basins
and the lithospheric condition under the Japanese arcs may be controlled by three factors.
These are: supply of sediments, horizontal stress, and isostatic balance. The supply of
sediments is estimated to be mainly controlled by the width and topographic relief of the
land area. The horizontal compressional stress would affect the island arc to make its
topographic relief steep and to increase the amplitude of tectonic deformation. It would
make the depressed area a deeper sedimentary trap and the uplifted area a higher source
of sediments. In the compressional stage such as the time interval I, the rate of sedimen-
tation should be high in the basins and coastal areas. This deduction agrees well with the
above-mentioned rate of sedimentation. After the detachment of a dense lithospheric
slab under an island arc system, isostatic rebound would be expected, and the land area
should be more extensive and the amplitude of topographic relief should decrease, due to
decrease of horizontal stress. The sedimentation rate should be made lower. The ex-
tremely low rate of sedimentation in time interval II can be explained by this process.
According to FORSYTH and UYEDA (1975), the motive force of plate motion is mainly
caused by the subsidence of an oceanic lithospheric slab. If we follow this scheme, we
Magnetic Stratigraphy of the Japanese Neogene and the Development of the Island Arcs S 377
should find the influence of cyclic evolutionary trends of plate subduction on the plate
motion and ocean floor spreading. We can monitor the motion of oceanic plates by the
position of volcanic islands created by the activity of hot spots. In the Hawaiian hot spot
chain, the rate of volcanic progression has been smoothly accelerating in the last 4my
(JACKSON, 1976). The acceleration is consistent with the existence of Northeastern Japan-
Izu-Bonin-Mariana Arcs in which the inter-plate friction decreases with the evolutionary
trend.
The spreading feature at the mid-oceanic ridge can be demonstrated by the ocean
magnetic anomaly pattern correlating with the geomagnetic polarity reversal history.
In the East Pacific area where the Pacific plate is being created, a fossil spreading center
has been found by HERRON (1972). Because it is estimated that this center stopped spread-
ing at 10mybp (corresponding to anomaly 5), the East Pacific region must have had
two spreading centers prior to 10mybp. This age corresponds to the boundary between
time interval II and III, when the new cycle of subduction started. Because the early
stage of the evolutionary cycle can be compared to the Chile type subduction in which
inter-plate coupling is strongest, the pulling force for plate motion should have been de-
creased at the subduction area. The decrease in the pull of the Pacific plate was able to
close the fossil spreading center and to decrease the total spreading rate.
In conclusion, the rate of sedimentation deduced from magnetostratigraphy is a very
useful parameter in reconstructing past tectonic conditions. Using the sedimentation rate
data, the Japanese Neogene and Quaternary can be divided into three major time inter-
vals. The features of tectonic evolution and paleogeography in each time interval are
consistent with the "cyclic evolutionary model," originally proposed by KANAMORI (1971,
1977) to explain the process of oceanic lithospheric subduction. The pulling force of plate
motion at the subduction area, represented by FORSYTH and UYEDA (1975), seems to con-
trol the Pacific plate motion because of the correspondence between the plate motion and
cyclic evolutionary trend of plate subduction.
Grateful thanks are expressed to Professors Hakuyu Okada, Shizuo Yoshida of Shizuoka University, and
Professor John H. McD, Whitaker of Leicester University for discussion and critical reading the manuscript,
and to Professors Nobu Kitamura, Hisao Nakagawa, and Drs. Toyosaburo Sakai, Kenshiro Otsuki of Tohoku
University, and Professor Asahiko Taira of Kochi University for valuable suggestions. Acknowledgements
are also due to Dr. Kenichi Manabe of Fukushima University and Mr. Kazuo Yoshida of Tohoku University
for discussion and magnetostratigraphic information in the Sendai and Boso areas.
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FORSYTH, D. and S. UYEDA, On the relative importance of driving forces of plate motion, Geophys. J.R. Astr.
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83, 1671-1692, 1972.
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ISHIDA, M., Seismicity and travel-time anomaly in and around Japan, Bull. Earthq. Res. Inst. Tokyo Univ., 48,
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KIMURA, K., Magnetic stratigraphy of the Late Cenozoic sedimentary sections in Boso Peninsula, Niigata area,
and Oga Peninsula, Japan, J. Geol. Soc. Jpn., 80, 579-592, 1974.
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and Upper Mantle of Japanese Area (Part II), edited by M. Gorai and S.Igi, pp. 38-60, Geol. Surv. Jpn.,
Kawasaki, 1973.
KOBAYASHI, K. and N. ISEZAKI, Magnetic anomalies in Japan Sea and Shikoku Basin and their possible tectonic
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KOBAYASHI, Y. and K. NAKAMURA, Restoration of tectonic stress field of Tertiary Southwest Japan by means
of dikes, Abstr. Int. Geodyn. Conf., pp. 86-87, 1978.
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NAKAGAWA, H., N. NIITSUMA, and I. HAYASAKA, Late Cenozoic geomagnetic chronology of the Boso Peninsula,
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NAKAGAWA, H., N. NIITSUMA, K. KIMURA, and T. SAKAI, Magnetic stratigraphy of Late Cenozoic stages in
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Burckle, Micropaleont. Press, pp. 64-70, New York, 1975.
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Japan, Quatern. Res., 7, 294-301, 1977.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 99, NO. Bll, PAGES 22,165-22,185, NOVEMBER 10, 1994
Aeromagnetic evidence for a buried Early Cretaceous
magmatic arc, northeast Japan
Carol Finn
U.S. Geological Survey, Denver, Colorado
Abstract. Positive aeromagnetic anomalies, recent drilling, and models constructed
from these data delineate the plutonic roots of the Early Cretaceous Kitakami
magmatic arc in northeast Japan. Buried plutons, mostly offshore, produce belts of
positive magnetic anomalies. These anomalies and magnetotelluric data suggest that the
plutons form a batholith 70-120 km wide, nearly 800 km long, and 10-15 km thick. The
batholith may mark the location of the main Kitakami arc. Most of the exposed
Kitakami plutons are 2-20 km in diameter; some are 3 km thick. The small plutons line
up along NW trending faults; some may have been satellite vents that tapped into the
magma supply of the main arc. The batholithic roots of the main arc now compose
almost half of the modern Japan forearc basement. Steep magnetic gradients, offset
anomalies, and basin stratigraphy portray extensive faulting of the Kitakami batholith
during oblique subduction in the Late Cretaceous and rifting in the Miocene. The
eastern boundary of the Kitakami batholith lies between 90 and 140 km west of the
modern trench, much closer than the 300-km distance between the active arc and
trench. The Early Cretaceous forearc basin and accretionary prism may underlie the
modern forearc basin east of the batholith, but clear evidence is lacking. Much of the
Early Cretaceous margin, including most of its forearc therefore is missing. How the
material was removed is unknown: it could have been strike-slip faulted, eroded by
subduction-related processes, or both.
Introduction
Northeast Japan (Figure 1) has been a convergent margin
for much of its history since the Mesozoic. Exposed rocks,
faults, and structure hint at past subduction-related pro-
cesses and have led to many models of Mesozoic and
Cenozoic tectonic events in northeast Japan [Uyeda and
Miyashiro, 1974; Dickinson, 1978; Kimura et al., 1983;
Kimura, 1985; Maruyama and Seno, 1986; Jolivet et al.,
1991; Kimura, this issue]. Mesozoic rocks and faults are
buffed beneath younger rocks, vegetation, and water. In
order to constrain tectonic reconstructions and the modern
forearc basement composition of northeast Japan, more
information on its Mesozoic structure is necessary. Distinct
aeromagnetic anomalies are associated with exposed and
buried rocks of an Early Cretaceous arc-trench system and
thus provide the best means for delineating basement com-
position in the region. Other data sets for the region are not
as effective in finding buffed crystalline rocks [Finn, 1994;
Zhao et al., 1992; Zhao and Hasegawa, 1993]. The major
objective of this paper is to demonstrate that the aeromag-
netic maps and 2-dimensional models across the margin delineate the mostly buffed batholithic (?) roots of the Early
Cretaceous magmatic arc.
Plate Tectonic History
In the Late Jurassic to Early Cretaceous, an accrctionary
prism and a collided continental block formed part of the
This paper is not subject to U.S. copyfight. Published in 1994 by the
American Geophysical Union.
Paper number 94JB00855. eastern margin of Asia of which Japan was a part. In the
Early Cretaceous, fast (more than 20 cm/yr), oblique sub-
duction of the Izanagi plate [Engebretson et al., 1985]
produced the Kitakami continental magmatic arc (Figures 2
and 3). Volcano-plutonic activity in the Kitakami arc ceased
by about 110 Ma [Shibata, 1968] but continued along the
Asian margin through the Paleogene [Taira et al., 1983;
Parlenov et al., 1978]. The Rebun-Kabato tholeiitic subma-
fine island arc (Figures 2 and 3) was interpreted to lie
trenchward of the continental arc [Niida and Kito, 1986].
Deposition of forearc basin sediments, the Yezo Group,
on oceanic mafic basement and accretionary prism of the
Sorachi Group (Figures 2 and 3) occurred throughout the
Cretaceous [Okada, 1983]. The accretionary prism grew
eastward through the Paleogene [Kimura, this issue]. The
subducted accretionary prism of the Sorachi Group was
serpentinized and metamorphosed under blueschist facies
P/T conditions. Episodic, rapid underplating of accreted
sediments may have caused uplift of the metamorphic rocks
within the Yezo forearc sediments during the Cretaceous
(Figure 2) [Kimura, this issue]. These uplifted rocks now
compose the Kamuikotan metamorphic belt (Figures 2 and
3) [Niida and Kito, 1986].
Opening of the Japan Sea (Figure 1) in the early Miocene
[Otofuji et al., 1985] (see Tamaki [1988] and Jolivet et al.
[this issue] for reviews) rifted Japan from the Asiatic margin.
Opening of the Kuril basin (Figure 1) at the same time
induced dextral strike-slip faulting in Hokkaido and Sakhalin
(Figure 1) [Kimura et al., 1983; Jolivet, 1986]. Silicic "Green
Tuff" volcanism accompanied both rifting events and now
covers many of the older rocks in northern Honshu and
much of Hokkaido [Togashi, 1983; Nohda et al., 1988;
22,165
22,166 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
SIKHOTE ALIN SAKHALIN
JAPAN SEA
PACIFIC OCEAN
Figure 1. 44'
42'
40'
38'
36' 140' Rebun 142'
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:.:::.::.::::::::: "::::::.:: ::-:::: "!::::::::: ..... :::::::::::::::::::::: ..... ::.'-'-':.F.::-'-.--::- '.::it-'.'-;::: ........... ::::::::::::::::::::::::::::::::::::::: i ;-:,:i,:::,ii::ili.:. :::::::::::::::::::::::::::::::::::::::::::::::::::::::::: ::!::i:i!iii!i.:. i:;:: i?. ,:.i ::::!::?:!;!;' :' H" :O.KA ! :B:O..::i::.::::...: ' ::':,i::?iiii::i::iiiiii iii[::;::/:!ii::i:.f:::: :::!!!::i::5i[:5:::.'..'i[::::5'"i..".-i::: :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: '.:::.::[:i!i:;.;i::i::::::!fi::i:'is."! .::::::ii:::-*-:.'...:'iii::i."!!: ;h........'.'::. ""::':.i::::::::: ::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: :!"::.:.::!:.:!:!: .... .'-'i.::..s,:.::'.:.-::......[:i::..:','-::i:!::.-$:' ::::;.':'.::
.- ..... -:::.-- .............. ' ........ M.is ..............................................................
"; ::::::::::::::::::::::::::::::::: ......... ,.:.::-e*..-.'.--! .. .'';:' ;':':---:"----":.....!:,..-i- ....... . ii??i':;;:-:';:?: ........... :' ' ':' '" ":jliiii:: :;;i;;iiii117!?':':""'::j:½J:' :-:. ß .... - :::!!!!i.'!"::';:...: ..... ::!i.::i!:'.i::.,:i.-':'::4ii:,. '?:.::!':'---.:':::'.-ii'" :-"'::":: ....... :::*:****':':':*:"..': ::..:i::ii::::::::i:;;*4i:.::`,.iiii!i: ..--".:-:5"::'":t.i!.'. ø' ,:.;, .... ."...
:-'. ?.*'i;:;**,?:.;;t:,:/:':.:";:i%¾:*"";"'"'? '*'". ..................... : .... "'::.:::*¾%i;;'."'"'":*;:i;:;i::*:*-'"' ' . :? .Z:$::R4.!::..? ' ..::-:.:.::::i:i:::...'::.. :::?,.::! ======================================== .?: -'.E::::.,...:, ..... :.:.; ..::::!'ii ::
:`5!!!!:::!!!ii`:i::i*:::::::::;:::::;:``:*i::i:::::: i:,:::!* ........... :ii ...... !j]`t!L::i*::::!U:L:*:ii:*:::::.::i::::.{;:::;:!::::;!:::::!:::::q::i!i:;ii:i '
i,c' .::::"":. ,. `:::41:::1?:`ii!i?:cjr*::ii;ii:::::::::::::*iiii:!i:i!:::!i*;;iiii 40'
-i:ii:;'.": .::.::;S::.,::'i.:.:.. '":':i:.;.:',;" :;.... ':ii:i:::iii::..</:::ii,::ii:::i::ii:.'::.:..!!::!::!i:;::
,:ii"":':- -*:':i!:::**: *!:._:...-:;iiii.. ..' ...... ;...!i..L*.:.:';J;';;,:i:::i:i:.' ."
(D:":-!;'"$ii:ij',iiD?? :t%..:: ::]i: .............. ';;*;:: ...................
:..:.i...f :::.:!!..4....... :: m::!::.,; ...... ::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: '::::"::::::' ....... * ..........
::.i** ::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: ". .......
-!-' - ':::,::;i::;:!:.::i!i:::::i! i:.i:.*!:: ::::::::::::::::::::::::::'.:% 3 6'
ii::ii::::::i ..... ===================================================================================== :::::::::::::::::::::::::::: "::.;: ::- ...... !::i::/D::::::::!::::::i::i*::*::::*:::i:::i*:::::::::::;f:.::.> '
'..::,-. .,:,:,.:::**f:11!?:iii;:i!::iiii?:i::?:::iiiiii:,i:.:.::;::.:i;:?:i:::" .............. :.'"
140 ø 142 ø
o lOO 2oo KM
Topographic and location maps of the study area. Meters
ß -- . 1000
i.. ! 8oo
"*"""':'"':"' '"" '" 600
4OO
....... . '":.:.::ia 0
-1000
-2000
-3000
-4000
Geological Survey of Japan, 1992]. Initiation of subduction
of the Pacific plate at the Japan trench in the late Miocene
was followed by construction of the modern Quaternary
volcanic arc (Figure 1) [Honza, 1980]. Collision of the Kuril
arc (Figure 1) in the Pliocene uplifted the Hidaka Mountains (Figure 1) [Komatsu et al., 1983, 1989; Kimura et al., 1983].
Pliocene to Quaternary uplift of the Kitakami Mountains
[Kobayashi, 1941; Chinzei, 1966] and Abukuma Highland
(Figure 1) [Kubo and Yamamoto, 1990] exhumed some of
the pre-Tertiary basement, but much of it, particularly
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,167
CRETACEOUS ARC-TRENCH SYSTEM
NORTHEAST JAPAN
Kitakami continental (100 MA) magmatic arc
0 - ;.: ;.: q.: ;.: ;.:.?:......-....._..._... -..- . -..- ... -.-.- -. -.-. - -. - -. - -. - -. -......- .-. . '"]','-.',".',"-','-'-":.":,':'.'",':','"-': V&.afoa3cb'si".':',':'.":,':',':',':.",':.".'::" Trench axis ] 0 . . ......... :.? :.: :.: :.? :; ß .... :....:..:..:..:..:..'. :.:. :... :........,.:.... . -J-
t/ '+'+'+ato xxt_ ___ _x/++ + Sorachiforearcbasement ..... :'"':':"'"' :i'."'.':'.:!i [ ...... ß . .r .... + + + + .:.:.:.:,:.:.:.:.:.:.:.:.:.:,:.:.:,:.:.:.:,
, ++ ?9r+a .chiq ec ...... '-' '"'"'"' '-'"":::i:... ., . .... . .......'iii:::'::i:!- 10
15 iiiiii-- 15
20 ................................................................. /.t..,....::.:.-.., ...................... ................................................................................................................................................ ii::iii' 0
30 30
0 50 100 KILOMETERS
I , [ [ , I [ , I
Vertical exaggeration = 200 percent
Figure 2. Cartoon of a section across the Early Cretaceous convergent margin (modified from Niida and
Kito [1986]).
offshore and in north central and northwestern Honshu
(Figure 1), remains covered by Cenozoic volcanic and
sedimentary rocks (Figure 3).
Aeromagnetic Data and Interpretation
Data
The aeromagnetic survey covering the land area in Hon-
shu and western Hokkaido was funded by the New Energy
Development Organization, a Japanese governmental
agency, and flown at a barometric altitude of 1372 m along
northwest and east trending lines at 3-kin intervals. From
central Hokkaido to the east, many surveys were flown by
the Geological Survey of Japan at various barometric eleva-
tions and line spacings [Makino et al., 1992b]. Several
surveys were flown over the marine areas at a barometric
elevation of 457 m along northeast trending lines spaced 3-4
km apart [Makino et al., 1992b]. Corrections for the diurnal
variation of the magnetic field and the International Geomag-
netic Reference Field (IGRF) updated to the years of the
surveys were subtracted from the data [Makino et al.,
1992b]. I continued all data to an elevation 304 m above the
terrain using the method of Cordell and Grauch [1985] and
merged them into a unified map (Plate 1).
The most prominent feature of the aeromagnetic map
(Plate 1) is the Kitakami magnetic anomaly belt (KMB,
Figure 4) [Makino et al., 1992b], a north trending 50-km-
wide, 500-km-long belt of magnetic highs with amplitudes of
0-900 nT peak to trough. Farther south is the Jouban
magnetic anomaly belt (Plate 1 and JMB, Figure 4), a
300-km, north trending belt of magnetic highs with similar
amplitudes and wavelengths to those in the KMB. East of
and parallel to the KMB are several discrete, positive
anomalies in the eastern magnetic anomaly belt (Plate 1 and
EMB, Figure 4). Serpentinites of the Kamuikotan metamor-
phic belt (Figure 3) are associated with a 50-km-wide,
400-km-long belt of magnetic highs (Plate 1 and KSMB,
Figure 4) with amplitudes of 0-1500 nT peak to trough. I
make the case in this paper that these anomalies are related
to rocks of the Cretaceous arc-trench system. Short-
wavelength positive magnetic anomalies along the western
and northeastern sides of the aeromagnetic map (Plate 1) are primarily caused by Miocene and Quaternary volcanic
rocks.
Previous Interpretations
Previous workers regarded the source of the KMB as
basalts, an ultramafic dike or a combination of both [Ogawa
and Suyama, 1975; Segawa and Tomoda, 1976; Segawa and
Futura, 1978; Makino et al., 1992b]. Ogawa and Suyama
[1975] speculated that the basic and ultrabasic rocks intruded
along a shear zone in the early Cretaceous as part of the
Mesozoic volcanic arc system now exposed in northeast
Japan. Segawa and Furuta [1978] related basalts found in
drill holes in the area of the KMB to the Sorachi Group.
Metadolerite in one drill hole was also correlated to volcanic
rocks exposed in the Kabato Mountains (Figure 1) [Takig-
ami, 1984]; the magnetic anomaly belt (Plate 1) was ascribed
to the Rebun-Kabato tholeiitic arc (Figures 2 and 3) [Takig-
ami, 1984; Ikeda and Komatsu, 1986].
Previously, the positive aeromagnetic anomalies were
modeled with 10- to 40-km-wide and 5- to 20-km-thick
sources with intensities of magnetization of 3-5 A/m [Ogawa
and Suyama, 1975; Segawa and Tomoda, 1976; Segawa and
Furuta, 1978]. Inclinations in the present Earth's field (53 ø)
but westerly declinations of 45øW (8øE is the present decli-
nation) consistent with Cretaceous paleomagnetic directions
were used in the models [Segawa and Tomoda, 1976;
Segawa and Furuta, 1978]. However, the drilled basalts
have low Q values (below 1 [Saito and Tanaka, 1975]). The
modeled and present Earth's field magnetization vectors are
within about 30 ø of one another and so can be considered
almost collinear [Bath, 1968]. Without knowing the structure
of a body, remanent direction cannot be determined. There-
fore remanence could have been neglected in the models.
Ogawa and Suyama [1975] interpreted that the source of
the KSMB was Kamuikotan serpentinites. von Huene et al.
[1982] suggested that the source of the EMB was Paleogene
arc volcanic rocks based on the discovery of Oligocene
dacite in a nearby Deep Sea Drilling Project (DSDP) hole.
However, as they themselves noted, a magnetic low lies
directly over the area drilled, so that it is unlikely that
Paleogene volcanic rocks are the source of the EMB.
These earlier interpretations did not have the benefit of the
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22,168 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
4 ø
44 ø
43 ø
41'
40 ø
39'
38'
37'
36'
35' 140 ø 141 ø
Kitakami
Province
North
Kitakami
, t South
_. Kitakami ! 142 ø 143 ø 144 ø
I
I I EXPLANATION
EARLY CRETACEOUS ROCKS
I Miyamori and Hayachine ultramafics
ARC ROCKS
.....:.,,.,... Kitakami plutons Rebun-Kabato volcanics
FOREARC BASIN ROCKS
:i.:..-',,,. Yezo sedimentary rocks Kamuikotan metamorphic rocks
ACCRETIONARY PRISM
i:._'-'. $orachi oceanic crustal rocks and melange, Hidaka melange
Hidaka western greenstone belt
LATE CRETACEOUS
ARC ROCKS
':.? Abukuma plutons
0 100 200 KILOMETERS
I I I
Figure 3. Simplified geologic map showing the location of provinces and rocks comprising the Creta-
ceous arc-trench system. Most of the areas shown in white are covered by Cenozoic rocks. The boundary
between the north and south Kitakami Provinces is marked by the west trending bold solid line. The thin
solid lines in the Kitakami Province mark major faults. The dashed lines mark the Tanakura and Hatagawa
fracture zones [after Geological Survey of Japan, 1992; Otsuki, 1992; Kirnura, this issue].
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,169
onland aeromagnetic data, new geologic mapping, rock
property measurements and drill hole information. The new
data warrant a reinterpretation of the sources of the positive
aeromagnetic anomaly belts (Plate 1 and Figure 4). In the
next sections, I present a detailed description of the associ-
ation of positive aeromagnetic anomalies with rocks that
composed the Early Cretaceous arc-trench system.
The Cretaceous Magmatic Arc
Kitakami basement rocks. The Kitakami province ex-
tends from northeast Honshu, north to the Oshima peninsula
in southwestern Hokkaido (Figures 1 and 3). The oldest
rocks of the province consist primarily of pre-Cretaceous
marine and continental sedimentary rocks, many of which
were part of the accretionary prism [Geological Survey of
Japan, 1992] along the Asian continental margin. Rocks of
the Jurassic accretionary prism now form the basement of
north Kitakami (Figure 3). A collided continental block
composed of Paleozoic granite and sedimentary rocks along
with marine sediments forms the basement of south Ki-
takami (Figure 3). Most of the basement rocks of north and
south Kitakami do not produce aeromagnetic anomalies.
The collision of the continental block with the accretionary
prism uplifted and serpentinized ultramafic rocks [Saito and
Hashimoto, 1982], now exposed in the central part of the
Kitakami Province as the Hayachine and Miyamori Com-
plexes [Ozawa, 1984] (Figure 3). These rocks produce pos-
itive magnetic anomalies with peak-to-trough amplitudes
greater than 800 nT [Okuma et al., 1992; Okuma, 1993] (MY,
HA Figure 5).
Exposed Kitakami plutons. The plutonic roots of the
former continental magmatic arc (Figures 2 and 3) intruded
the north and south Kitakami basement rocks in the Creta-
ceous [Kanisawa, 1990; Kanisawa et al., 1984]. The sources
of most positive aeromagnetic anomalies with amplitudes of
300-900 nT peak-to-trough (Figure 5) over the Kitakami
Mountains (Figure 1) are exposed Cretaceous plutons. Al-
though the plutons are zoned from quartz diorite to granite
[Kanisawa, 1990], most produce simple "bull's eye" anom-
alies. The largest exposed pluton, the Tono, causes a com-
posite positive magnetic anomaly (TO, Figure 5), possibly
because of variations in susceptibility and thickness. The 2-
to 20-km-diameter intrusions have low remanence, with Q
values generally less than 0.2 [Kawai et al., 1971; Murata et
al., 1991] but magnetic susceptibilities as high as 55 x 10 -3
SI units and an average of about 19 x 10 -3 SI units [Kanaya,
1974; Ishihara, 1979, 1990].
Magnetotelluric (MT) data [Ogawa, 1992] define the resis-
tivities and thicknesses of three Kitakami plutons along a
profile in the Kitakami Mountains (A-A', Figure 5). Accord-
ing to the MT model, the Himekami (HI, Figure 5) and
Sakainokami (SA, Figure 5) plutons are 3 km thick (Figure 6)
and the Miyako (MI, Figure 5) is 15 km thick (Figure 6). The
plutons are associated with positive magnetic anomalies
(Figure 6). The anomaly over the Miyako pluton (MI,
Figures 5 and 6) is part of the KMB (two other exposed
plutons produce anomalies within the KMB: Tanohata (TN,
Figure 5) and Hashigami (HG, Figure 5)).
Plutons with average susceptibilities less than 10 x 10 -3
SI units are not magnetic enough to produce positive aero-
magnetic anomalies at the elevation of the survey. Hitokabe
(HT, Figure 5) [Kanaya, 1974] and all Kitakami plutons in Honshu west of the Kitakami Mountains and Abukuma
Highland [Sasada, 1985; Kubo and Yamamoto, 1990] (Fig-
ure 1) and southwest Hokkaido (K. Kubo, personal commu-
nication, 1991) have susceptibilities too low to cause aero-
magnetic anomalies (Plate 1 and Figure 5).
Exposed Kitakami volcanic rocks. The Kitakami arc vol-
canic rocks, the Harachiyama Formation, range in compo-
sition from rhyolite to basalt; some were erupted offshore
[Kanisawa et al., 1984]. These rocks form small outcrops
between the Hatagawa and Futaba fracture zones [Ya-
mamoto et al., 1989], the southern Kitakami Mountains, and
along the northeast coast of Honshu [Geological Survey of
Japan, 1992]. Posteruption uplift of the Kitakami Mountains
(Figure 1)[Kobayashi, 1941; Chinzei, 1966] presumably led
to erosion of most of the Cretaceous volcanic rocks. Sparse
measurements of the Harachiyama Formation record pre-
dominantly low susceptibilities: 0.01 x 10 -3 to0.1 x 10 -3 SI
units, except for a few measurements with values of 10 -3 SI
units (S. Okuma, written communication, 1991), and low
remanent intensities of less than 0.1 A/m (T. Tosha, written
communication, 1991). The volcanic rocks were altered by
intrusion of the plutons, but sections have been observed
where the lower part is altered and the upper part is not.
Unaltered sections might produce magnetic anomalies if
sufficiently thick [Okuma, 1993]. Positive aeromagnetic
anomalies are associated with some exposures of the Harac-
hiyama Formation (part of anomaly VO, Figure 5).
Interpreted buried Kitakami magmatic arc rocks. I inter-
pret that most of the positive magnetic anomalies offshore
Honshu into central Hokkaido (Plate 1 and Figures 4 and 5)
are caused primarily by buried Kitakami plutons. In south-
ern Honshu, the wavelengths and amplitudes (150-700 nT
peak-to-trough) of aeromagnetic anomalies P1-P5 (Figure 7)
are similar to those associated with exposed Kitakami plu-
tons. The presence of sedimentary rocks belonging to the
southern Kitakami Province east of the Futaba fracture zone
(Figure 3) near anomaly P1 (Figure 7) indicates that the
basement belongs to the Kitakami Province; Kitakami plu-
tons would not be unexpected. Anomaly P2, connected to
P1, has the same wavelength and amplitude of anomaly P1.
A drill hole penetrated a Kitakami tonalite at 300 m depth
[Abe and Ishihara, 1985] beneath anomaly P3 (Figure 7).
Positive magnetic anomaly P4 (Figure 7) has been inter-
preted to be due to serpentinite based on geomagnetic
variation data [Seto and Kitamura, 1990]. All exposures of
serpentinite in northeast Japan are related to collision zones
or uplifted accretionary prism. No evidence for a collision
zone or accretionary complex exists in the area of anomaly
P4 (Figure 7). A more geologically reasonable source is a
buried Kitakami pluton. Miocene and Quaternary volcanic
rocks mapped in the area [Geological Survey of Japan, 1992]
over the proposed pluton may contribute to the observed
magnetic anomaly. I interpret that the high-amplitude posi-
tive anomaly P5 (Figure 7) is caused by both a partially
exposed Kitakami pluton and overlying Neogene volcanic
rocks.
An offshore drill hole (1, Figure 7) encountered Late
Cretaceous and Recent sedimentary rocks and the top of a
Kitakami pluton at a depth of 1830 m below sea level [Japan
Association of Natural Gas, 1986]. In two other holes (2 and
3, Figure 5), Harachiyama volcanic rocks at least 800 m thick
at depths of 2700 and 2200 m were found [yon Huene et al.,
1982]. Drilling in Hokkaido (4, Figure 8) penetrated a Ki-
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22,170 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
44'
42'
40'
38'
36' 140' 142'
8
44'
42'
40'
38' nT
5OO
450
400
350
3OO
25O
200
150
100
5O
+ 36'
140' 142'
0 100 200 KM
I
Plate 1. Color shaded relief aeromagnetic map of northeast Japan. Locations of Figures 5, 7, and 8 are
marked by boxes 5, 7, and 8, respectively.
takami pluton at a depth of 4645 m without overlying
Cretaceous volcanic rocks [Yufutsu Research Group of
JAPEX Sapporo et al., 1992]. These drill holes underlie parts
of the JMB and KMB (Figures 4 and 5).
In Hokkaido, just west of the KMB, Early Cretaceous
metadolerite was drilled at a depth of 4265 m in the Nanporo
drill hole (5, Figure 8) [Saito and Tanaka, 1975; Segawa and Oshima, 1975; Takigami, 1984]. The Cenozoic formations
overlying the igneous basement in both the Tomakomai (4,
Figure 8) and Nanporo (5, Figure 8) holes are virtually
identical (details are given by Segawa and Tomoda [1976]
and Yufutsu Research Group of JAPEX Sapporo et al.
[1992]). Relative to the overlying formations, the metadoler-
ite occupies the same position in the Nanporo hole (5, Figure
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,171
43 ø
42 ø
41'
40'
39 ø 140 ø 141ø 142 ø 143 ø 144 ø
KSMB
38 ø i
37' -
36' -
35ø I -
0 100 200 KILOMETERS
Figure 4. Magnetic belts described in the text. JMB,
Jouban magnetic belt; KMB, Kitakami magnetic belt; EMB,
eastern magnetic anomaly belt; KSMB, Kamuikotan setpen-
finite magnetic belt. 8) as does the Kitakami pluton in the Tomokomai hole (4,
Figure 8). This and the similarity of their ages suggest that
the metadolerite correlates with the Harachiyama formation.
Like the Tomakomai drill hole (4, Figure 8), two layers of
arc volcanic rocks were found in the Nanporo hole (5, Figure
8). Measured susceptibilities for the two volcanic layers in
the Nanporo hole are about 25 x 10 -3 SI units, and
remanent intensities are 0.44 A/m [Saito and Tanaka, 1975].
A 50-km-wide model that contained both volcanic layers
with assigned total magnetizations of 100 x 10 -3 A/m
(greater than measured) caused a 100-nT positive magnetic
anomaly, far less than the 800-nT anomalies observed near
the drill holes (4 and 5, Figure 8). Adding reasonable
thicknesses (less than 4 km) of Cretaceous metadolerite with
a total magnetization intensity of 100 x 10 -3 A/m (much
higher than measured) at the bottom of the model produces
anomalies with total amplitudes <200 nT. The volcanic
rocks therefore cannot be the primary sources of the mag-
netic highs.
Rebun-Kabato Belt. Two to three kilometers of Early
Cretaceous [Takigami, 1984] calc-alkaline to tholeiitic hyalo-
clastites interbedded with mudstone, acidic tuff and sand-
stone [Nagata et al., 1986] are exposed in the Kabato
Mountains (Figures 1 and 3). To the northwest, in Rebun
Island (Figure 1), 2- to 3-km-thick assemblages of Early
Cretaceous andesites are interbedded with calcareous mate-
rials (Figure 3) [Ikeda and Komatsu, 1986]. Volcanic rocks
north of Rebun Island (Figure 1) (in Moneron Island) corre-
late with the Rebun and Kabato rocks [Piskunov and Khved-
chuk, 1976]. Although the Kabato andesites are geochemi-
cally similar to the Harachiyama Formation in the Kitakami
Mountains [Kanisawa, 1990; Nagata et al., 1986], the Ka-
bato basalts and Rebun rocks are more tholeiitic [Ikeda and
Komatsu, 1986]. This difference in chemistry led Ikeda and
Komatsu [1986] and Niida and Kito [1986] to propose the
existence of a tholeiitic arc (Figures 2 and 3) located trench-
ward of the Kitakami magmatic arc. The KMB (Figure 4)
was ascribed to the Rebun-Kabato tholeiitic arc [Takigami,
1984; Ikeda and Komatsu, 1986]. However, instead of a
positive anomaly, the Kabato Mountains are associated with
a magnetic low (KM, Figure 8), indicating that the volcanic
rocks are not the cause of the KMB.
Abukuma Province. The Abukuma province lies south-
west of the Kitakami Province across the Hatagawa fracture
zone (Figure 3). Its southwestern boundary is the Tanakura
fracture zone (Figure 3). The province is underlain by
Paleozoic granites, high-temperature/low-pressure metamor-
phic rocks, serpentinites, and oceanic mafic, ultramafic, and
pelitic rocks [Faure et al., 1986; Hiroi and Kishi, 1989]
interpreted to be a microcontinent [Faure and Natal'in,
1992]. Late Cretaceous plutons (Figure 3) [Shibata and
Tanaka, 1987; Shibata and Uchiumi, 1983] intrude the older
rocks; these plutons represent the roots of a magmatic arc
[Kimura, this issue] formed as a result of subduction of the
Kula plate. The nonmagnetic plutons are distinguished by
ovoid aeromagnetic lows against a background of short-
wavelength magnetic highs (Plate 1) caused by the mafic and
ultramafic rocks and young volcanic rocks [Ogawa et al.,
1979]. The mafic and ultramafic rocks under positive anom-
aly A1 (Figure 7) were interpreted by Ogawa et al. [1979] to
belong to the Abukuma Province and have been mapped as
such [Faure et al., 1986]. The sedimentary rocks underlying
the area of anomaly A1 (Figure 7) have also been interpreted
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22,172 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
41' 142' ...........
41'
nT
600
500
40'
39' 39'
142' 400
300
200
lOO
-lOO
0 25 50 KM
i
Figure 5. Gray-scale subset of the aeromagnetic map shown in Plate l. The location of the figure is
marked by the white box 5 in Plate 1. Squares with numbers mark drill hole locations. Polygons describe
Cretaceous plutons and the Miyamori (MY) and Hayachine (HA) ultramarie rocks. VO, Harachiyama
volcanic rocks; HT, Hitokabe pluton; SE, Senmaya pluton; TO, Tono pluton; HI, Himekami pluton; SE,
Senmaya pluton; SA, Sakainokami pluton; MI, Miyako pluton; TN, Tanohata pluton; HG, Hashigami
pluton; KB, Kuji Basin; TF, Taro fault. The black and white lines marked TFE show the interpreted
projection of the Taro fault offshore. The white line marked KF shows the interpreted projected of the Kuji
basin fault. A-A' marks the location of the MT profile shown in Figure 6.
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,173
700 600 5oo
400'
300
200
oo
o
400
-200
RESISTIVITY,
(l'-m)
::::::::::::::::::::::::::::::::::::::::::::: 100 :::::::::::::::::::::::::::::::::::::::::::::
::::::::::::::::::::::::::::::::::::::::::::: ::::::::::::::::::::::::::::::::::::::::::::: -o
p...,
- 20
Figure 6. Kitakami MT model [after Ogawa, 1992]. HI, SA, and MI are the Himekami, Sakainokami,
and Miyako plutons, respectively. The location of the model and magnetic profile is marked as A-A' on
Figure 5.
to belong to the southern Kitakami Province [Geological
Survey of Japan, 1992], and the mafic and ultramafic rocks
have been correlated with the Hayachine belt in south
Kitakami [Otsuki, 1992]. It is therefore possible that the
source of anomaly A 1 belongs to the Kitakami Province. The
source of the offshore positive anomaly between anomalies
A1 and P2 (Figure 7) could be Abukuma or Kitakami mafic
and ultramafic rocks or Kitakami plutons. Evidence is insuf-
ficient to choose between these alternatives. Positive anom-
alies south of latitude 38 ø and east of longitude 141045 ' are
related to sources within the subducting Pacific plate [Finn,
1994].
The Kitakami and Abukuma Provinces were originally
apart [Taira et al., 1983]; Late Cretaceous sinistral and
Miocene [Sasada, 1985] dextral strike-slip movement along
the NNW trending Hatagawa fracture zone (Figure 3)juxta-
posed them [Kubo and Yamamoto, 1990; Sasada, 1985].
Cretaceous Forearc Rocks
The Cretaceous forearc basin and accretionary prism are
exposed across a 200-km-wide, 1400-km-long belt in central
Hokkaido (Figure 3) and Sakhalin [Kimura, this issue]. A
typical section within the forearc basin, from bottom to top,
contains (1) the Kamuikotan metamorphic belt consisting of blueschist facies, Upper Jurassic-Lower Cretaceous [Shiba-
kusa, 1974; Ishizuka et al., 1983; Sakakibara and Ota, this
issue], serpentinite melange, (2) the Sorachi ophiolite suite
with serpentinized ultramafic rocks, pillow lavas, and radi-
olarian cherts [Ishizuka, 1981; Kito et al., 1986], (3) Lower
Cretaceous calc-alkaline volcaniclastic rocks similar to those
found in the Kitakami and Kabato Mountains [Girard et al.,
1991; Kanie et al., 1981; Kiminami et al., 1985], and (4) the
mid to Upper Cretaceous Yezo Group composed of terrigi-
nous clastic marine sediments and interbedded ash layers
[Okada, 1983]. Kamuikotan metamorphic rocks form an
anticlinal core with the Sorachi mafic rocks, melange, and
Yezo sedimentary rocks outboard. Farther east is the Cre-
taceous to early Paleogene accretionary prism (Figure 3)
composed of several belts including the Sorachi Group
[Kirnura, this issue]:
Serpentinites in the Kamuikotan metamorphic belt (Figure
3) produce high-amplitude positive magnetic anomalies
(Plate 1 and Figure 8) [Ogawa and Suyama, 1975]. MT data
from southern [Ogawa et al., this issue] and central H0k-
kaido [Ogawa, 1992] delineate a 0.5- to 2-km thickness for
the highly conductive srpentinites. Simple magnetic models
across three sections of the Kamuikotan belt (not shown)
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22,174 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
38'
37' 141 ø 142 ø
-.. ' ..-'.'-:" "-"-'.::.: '.::,: -.,. '- i :i5i!:!:'--5- !.:..-.i .......... '..:!:--::i:" !:i!:'..!!::::....,.. .
' :?::::-:-i- : '"'.-'.'.. :. ..:'i '!!' .'.::.':-:::i:.i:..i .... i¾:#:- --'"---..-":::.:::.!:: ::::.
ß - ..---.. .............. :.-...-:..-,..k$.K:-'. .'..,:.-..x .' .'. - ...... -.--- ,,.-.,'-: --:
..:.!:i:ii...:..:.:.?..: ? :::- ??"--''"-:: ..:::,::::ii?iii!?i!?iiii?i?i? .::.:: ..... ;.!::!ii-iiii:::."
ß ":'[:5:.i: .....:.!.. ß -'-'.-'-'5...-': ...... :555:-"- .: . 5 :.'.'::::' .::::::':-:.. ? :::::z: . ::' 5..'..::::.::: :.: !!5 ':..'.::.: 5:-.:::::::. .v
-- .:.; "-.:??.i . ,.,.:.'.: . .:..
' :...... '¾
141' 142' 38'
37'
B' nT
0 25 50 KM
i
Figure 7. Gray-scale subset of the aeromagnetic map shown in Plate 1. The location of the figure is
marked by the white box 7 in Plate 1. Square with number marks a drill hole location. Polygons describe
Cretaceous plutons. White lines mark fault locations. Letters denote anomalies referred to in the text.
B-B' marks the location of the gravity and magnetic model shown in Figure 9. 350
300
250
200
150
100
50
demonstrate that 0.5- to 2-km-thick serpentine blocks have
susceptibilities that range from 75 to 125 x 10 -3 SI units,
similar to the 82 x 10 -3 SI unit average measured from the
Miyamori complex (MY, Figure 5) [Okuma et al., 1992]. The
MT [Ogawa et al., this issue] and magnetic data indicate that
the serpentinites broaden beneath the surface and are wider
than the outcrop patterns (Figure 8).
The alignment of the EMB with the KSMB (Plate 1 and
Figures 4 and 8) suggests a similar serpentinite source.
However, one of the anomalies in the belt (east of S', Figure
8) is also connected to the KMB (Plate 1 and Figures 4 and
8) and may be caused by plutons. The Sorachi drill hole on the east side of the KMB (6,
Figure 8) tells a different story than the nearby Nanporo drill
hole (5, Figure 8). Metabasalt was drilled at a depth of 3705
m. Its measured susceptibility is 75 x 10 -3 SI units with a
remanent intensity of 0.8 A/m [Segawa and Furuta, 1978].
The overlying sedimentary section contains Cretaceous
forearc basin rocks, members of the Yezo Group and Ceno-
zoic sedimentary rocks, shuffled together because of the
Kuril collision.
Farther north, the Rumoi drill hole (7, Figure 8) encoun-
tered part of the middle Yezo Group overlying Early Creta-
ceous andesite correlated to that exposed in the Kabato
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,175
45'
44'
43'
42'
41' 140' 141' 142' 143'
q- q- -!--
q- q-
140' 141ø 142' 143' 45'
44'
43'
42'
41' nT
0 50 100 KM
Figure 8. Gray-scale subset of the aeromagnetic map shown in Plate 1. The location of the figure is
marked by the white box 8 in Plate 1. Squares with numbers mark drill hole locations. Polygons describe
Cretaceous plutons in Honshu and southwest Hokkaido; Early Cretaceous volcanic rocks in the Kabato
Mountain (KM); Kamuikotan serpentinites in central Hokkaido. Thrust faults are marked by lines with
teeth on the top plate; locations are from Miyasaka and Mitsui [1986]. KSMB, Kamuikotan serpentinite
magnetic belt. C-C' marks the location of the Shimokita modeled profile (Figure 10). S-S' shows the
location of a seismic refraction profile. 300
200
lOO
-lOO
-200
-300
-400
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22,176 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
Mountains [Japan Association of Natural Gas, 1992]. The
Enbetsu drill hole (8, Figure 8) detected duplexed sections of
the Cretaceous forearc sediments of the Yezo Group like the
Sorachi hole but no Early Cretaceous volcanic rocks. A
low-amplitude magnetic high in the Rumoi area (7, Figure 8)
is probably associated with nearby and offshore, highly
magnetic Miocene basalts, not the Early Cretaceous andes-
ite. No magnetic highs exist over the Enbetsu area (8, Figure
8). The presence of the Yezo forearc sediments in these
areas suggests that the Kitakami plutons or volcanic rocks
are not present.
Cretaceous Accretionary Prism
Eastern Hokkaido has low magnetic relief (Plate 1), re-
flecting the nonmagnetic Crctac½ous-Pal½ogcn½ accrctionar¾
prism rocks (Figure 3). High-frequency magnetic anomalies
in the area are produced by Miocene and Quaternary volca-
nic rocks.
Summary
The association of positive aeromagnetic anomalies on
land with exposed and drilled plutons (Figure 5) suggests
that the JMB and KMB are primarily due to the Kitakami
plutons with possible contributions from volcanic rocks of
the Harachiyama Formation. This interpretation differs from
previous models of the source of the KMB that suggest only
mafic source rocks [Ogawa and Suyarna, 1975; Segawa and
Oshirna, 1975; Segawa and Tornoda, 1976]. The following
model experiments test the hypothesis that buried Kitakami
plutons cause the KMB and JMB and investigate whether
the EMB is caused by serpentinites or Kitakami plutons.
Modeling
Models of the geometry and rock properties of sources
that produce magnetic and gravity anomalies guide us to-
ward geologic explanations for a study area. The models
usually only constrain a range of possible sources for anom-
alies and do not depict definitive structures and rock types. 1 ß
For two areas, I applied a 2-dmensonal, forward-and- inverse modeling program [Webring, 1985] to the magnetic
and gravity data. Two profiles across the southern (Jouban,
Figure 1) and northern (Shimokita, Figure 1) coasts and
offshore areas were chosen to delineate the buried Creta-
ceous arc-trench system.
The same procedure was used for all models discussed in
the text. Incorporation of all available information con-
strained starting models specified by body corners that
describe vertical blocks and magnetization and density con-
trasts. Perpendicular to the profile, the blocks were extended
distances usually of the order of 100-200 km determined
from the geologic, topographic, magnetic, and gravity maps
[Kornazawa et al., 1992; Kono and Furuse, 1989]. To
eliminate end effects parallel to the profiles, each starting
model was continued with the layers at the ends of the
section out to -1000 km. The program adjusted the starting
model so that its magnetic and gravity attraction fit the
profiles of observed data. I controlled the evolution of the
final model by allowing only a few parameters to vary in each
modeling attempt and constraining those parameters to a
specified range.
In all cases, magnetizations of modeled bodies were
assumed to be in the present Earth's field direction of
inclination 50 ø, declination -6.5 ø, and intensity of 46,000 nT [Geological Survey Institute, 1983]; remanent magnetiza-
tions are low relative to the susceptibilities. Initial suscepti-
bility values were based on available rock property measure-
ments [Murata et al., 1991; Tanaka and Kanaya, 1986, 1987;
Kanaya, 1974].
The models contain the subducting Pacific plate (Figure
9a; not shown in the other models). The geometry of the
subducting Pacific plate was not allowed to change in the
modeling and was included only for completeness. The
location and structure of the plate (Figure 9a) were inferred
from seismic refraction data [Nishizawa and Suyehiro, 1990;
Suyehiro et al., 1990; yon Huene et al., 1982] and from the
assumption that Benioff zone earthquakes [Nishizawa and
Suyehiro, 1990] lie within the crust and mantle of the
subducting plate. The densities [Nafe and Drake, 1957] were
estimated from seismic velocities [Suyehiro et al., 1990; yon
Huene et al., 1982]. The crust was assigned a single normal
magnetization, rather than variable values reflecting normal
and reversed rocks. This assumption is reasonable because
the observed magnetic field in the area is not sensitive to the
deep subducting crust. Crust and mantle densities of the
lower overriding plate were calculated [Nafe and Drake,
1957] from seismic velocities [Nishizawa and $uyehiro,
1990; Suyehiro et al., 1990; Ludwig et al., 1966; Asano et al.,
1981; yon Huene et al., 1982]. These are not well con-
strained. The flatness of the boundary between the upper
and lower crust at 16-km depth is arbitrary (Figures 9 and
10); no constraints on its shape are available.
Marine gravity and magnetic data [Makino et al., 1992a;
Nishimura and Murakami, 1977; Nishimura et al., 1984]
were used to model part of profile B-B' (Figure 7) and all of
C-C' (Figure 8). Few other data were available to constrain
the models; only general information can be obtained from
them. The assigned 16-km depth for the bottom of the
plutons was extrapolated from the MT data over the Miyako
pluton (Figure 6) [Ogawa, 1992] and the eastern Abukuma
Highland (Figure 1) [Ogawa, 1992]. Part of the Jouban (B-B',
Figure 7 and Figure 9) model was constrained by seismic
refraction [Suyehiro et al., 1984] and reflection data [Sakurai
et al., 1981]. Data from a north trending reversed seismic
refraction line (S-S', Figure 8) [Asano et al., 1979] con-
strained the location of two layers of the Shimokita model
(C-C', Figure 8 and Figure 10).
Kitakami plutons with susceptibilities ranging from 34.4 to
39 x 10 -3 SI units can account for most of the observed
positive aeromagnetic anomalies (Figures 9 and 10). Because
the geometries of the plutons are not well-constrained these
values can only be considered gross estimates (Figures 9 and
10). The plutons do not produce significant gravity anomalies
(Figures 9 and 10) [Tanaka and Kanaya, 1986, 1987]. The
modeled densities for the plutons of 2660-2680 kg/m 3 fall
within the measured range [Tanaka and Kanaya, 1986, 1987]
and in the Shimokita model (Figure 10) correspond to a 6.2
km/s seismic refraction layer [Asano et al., 1979]. The
widths of the modeled plutons are about 70 km in one of the
Shimokita models (Figure 10b), 90 km in the Jouban (Figure
9), and 120 km in the other Shimokita model (Figure 10a).
Near the coast, the tops of the plutons are about 600 m and
2000 m deep in the Jouban (Figure 9) and Shimokita (Figure
10) models, respectively. The average depth to the top of the
pluton farther east for all models is about 5 km. The low
densities of rocks (2100-2230 kg/m 3, Figures 9 and 10)
beneath the ocean (1030 kg/m 3) correspond to seismic re-
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,177
A)
tO 600
400 < o
-2oo
200 o -lOO
o
50 B B'
1030 Trench
3010 Continental lower crust Oceanic cxust
2810 (52.5) 254O
Oceanic mantle Oceanic asthenosphere 3210 3220 Continental mantle 3240
0 20 40 60 80 100 120 140 160 180
DISTANCE (km)
B) 2OO
600
4OO
oo o
-200
450 300 JMB
EXPLANATION
EARLY CRETACEOUS
[ Buried Kitakami plutons
Exposed Kitakami plutons
CRETACEOUS (?)
'- '."., S edim e n ta ryr o c ks LATE CRETACEOUS
;;..': Abukuma plutons
t::ii':::i';:{ Sedimentary rocks
CENOZOIC
Sedimentary rocks ø Observed
Calculated
Figure 9. Results of the magnetic and gravity modeling of the Jouban profile (B-B', Figure 7). Numbers
on bodies denote densities in kilograms per cubic meter, followed by susceptibilities x 10 -3 SI units
greater than 0 x 10 -3 SI units in parentheses. No vertical exaggeration. (a) Complete model including the
location of the subducting plate. (b) Closeup of western half of model in Figure 9a. JMB indicates that the
positive anomalies are part of the Jouban magnetic belt. Thick horizontal lines delineate horizons taken
from a north trending refraction model projected from 45 km to the south [Suyehiro et al., 1984].
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22,178 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
A C EMB C' 300 I._ I .... I .... I .... I .... I 'o ' ' ' I .... _
150
0
-45O
o < -100 .............. - '--- I I , , , , I , , , , i , , , , I , , , , I , , , I -
o ........................................................ .o3.o ..................
,--...:-:-:.::::::::::::::.::::: ß , ß,
0 .-'; k ',."." m . ............. ' ::''.:i." ,, ,, 2720-
20 i .... i .... i , ,3010 , ! .... i .... .... .
300-- I .... I .... i .... i .... I .... I .... _ _
,o o
-150 -300
-450
20 1030
,..... ............. _ , '"':' ::.!:;i..:i/:i':. ::...:. ß . ....... ;. :._.......:-,' :.:, ................. ' '."" '.' Z,?" 2730 2625(116) .... ,, ,3top, , .... ,',,,, I,,,, ,,,, 0 25 50 75 100 125 150
DISTANCE (kin)
'"'"' ...... Jurrassic (?) sedimentary rocks
EARLY CRETACEOUS ROCKS
Buried Kitakami plutons
'.'? Yezo sedimentary rocks Kamuikotan metamorphic rocks
basement EXPLANATION
:'" '.,?: Late Cretaceous sedimentary rocks
Cenozoic sedimentary rocks
ø Observed
Calculated
Figure 10. Results of the magnetic and gravity modeling of the Shimokita profile (C-C', Figure 8). (a)
Pluton model. KMB and EMB indicate that the positive anomalies are part of the Kitakami and eastern
magnetic belts, respectively. (b) Serpentinite model. Thick horizontal lines delineate horizons taken from
the refraction model [Asano et al., 1979].
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,179
fraction velocities of 1.7-3.3 km/s [Asano et al., 1979;
Suyehiro et al., 1984]. Cores from the top 1500 m of offshore
drill holes (2, 3, and 9 in Figures 5 and 8) [von Huene et al.,
1982] and interpretation of seismic reflection data from the
entire offshore area show these rocks to be Neogene to
Quaternary marine sediments [Honza et al., 1978; Commit-
tee for Co-ordination of Joint Prospecting for Mineral Re-
sources in Asian Offshore Areas (CCOP), 1991].
The densities of the layers beneath the Neogene to Qua-
ternary sediments in the models (2520-2660 kg/m 3 , Figures 9
and 10) were calculated from seismic refraction velocities
[Nafe and Drake, 1957] of 4.7-5 km/s from data in the area
of both the Jouban [Suyehiro et al., 1984] and Shimokita
(S-S', Figure 8) [Asano et al., 1979] models. The offshore
drill holes (1 (Figure 7), 2, and 3 (Figure 5)) contain Late
Cretaceous (belonging to the Kuji and Miyako, not Yezo,
Groups) and Paleogene marine sedimentary rocks [Japan
Association of Natural Gas, 1986; von Huene et al., 1982]
that probably correspond to the layers.
The models have different bodies east of the plutons. In
the Jouban model (Figure 9b) a 5.5-5.7 km/s layer corre-
sponds to a modeled density of 2740 kg/m 3 [Suyehiro et al.,
1984, 1990]. The calculated susceptibility for this body is
zero, indicating sedimentary rather than crystalline rock
types. The velocity and density are too high for the sedi-
ments to belong to the Yezo Group but may be appropriate
for the Sorachi Group or the Kitakami Jurassic accretionary
prism. The data are not sufficient to determine the source
however.
West of the Kitakami plutons in the Jouban model is the
Abukuma batholith (Figure 9b). The high-density 2770-2800
kg/m 3 bodies on the west end of the Shimokita models
(Figure 10) probably represent the Jurassic accretionary
prism.
The eastern positive magnetic anomaly, part of the EMB,
in the Shimokita model is interpreted in Figure 10a as a
Kitakami pluton but also can be modeled by a 45-km-wide
Kamuikotan serpentinite body (Figure 10b). The 2-km thick-
ness of the proposed serpentinite body and the structure of
the adjacent rocks were constrained by MT [Ogawa et al.,
this issue] and seismic refraction data from Hokkaido [Fujii
and Moriya, 1983], respectively, and extrapolated to the
offshore area. The calculated susceptibility of 116 x 10 -3 SI
units is within the 75-126 x 10 -3 SI unit range, and the width
is within the 20- to 50-km range modeled for exposed
serpentinites in Hokkaido.
Kamuikotan serpentinites are exposed in Hokkaido (Fig-
ure 3) as the cores of anticlines in-board of the Yezo and
Sorachi Formations [Kimura, this issue]. If Kamuikotan
serpentinite is the source of the EMB, the 2520 and 2730
kg/m 3 layers (Figure 10b) would most likely represent sedi-
mentary rocks of the Yezo and mafic rocks of the Sorachi
Formations, respectively.
Available data are insufficient to unequivocally determine
the source of the EMB. However, inferences can be made
that the most likely source of the EMB is serpentinite. The
wavelength of the anomalies within the ESB are more similar
to those in the KSMB than to those in the KMB (Plate 1 and
Figure 4). If Kitakami plutons underlie the ESB, the main
arc would have been 150 km wide and plutonism would have
occurred over a distance of 300 km (including the plutons in
southwest Hokkaido). These widths are greater than ob- served for modern arcs and are especially wide, considering
the short duration of plutonism in the Early Cretaceous.
Discussion
Cretaceous Magmatic Arc
Kitakami batholith. Drilling, correlation with exposed
rocks, and models (Figures 9 and 10) demonstrate that mag-
netic belts KMB, JMB and possibly the EMB (Figure 4) are
primarily due to Kitakami plutons. Harachiyama volcanic
rocks may overlie the plutons and contribute to the observed
magnetic anomalies, particularly for the Shimokita area, where
drilling (2 and 3, Figure 5) encountered Cretaceous andesites.
However, other onshore (near anomalies P1, P3 (Figure 7) and
4 (Figure 8) [Abe and Ishihara, 1985; Yanagisawa et al., 1989;
Yufutsu Research Group of JAPEX Sapporo et al., 1992] and
offshore drill holes (1, Figure 7) in areas associated with
magnetic highs [Japan Association of Natural Gas, 1986]
encountered plutons but no Cretaceous volcanic rocks.
The MT data [Ogawa, 1992] show that exposed plutons
under the KMB and JMB are about 10-15 km thick (Figure 6).
The seismic refraction data offshore [Asano et al., 1979] could
be interpreted to show that the buried plutons may be nearly as
thick, but this is speculative. The relatively large thickness and
area corresponding to the magnetic belts suggest that the
plutons are part of a batholith. According to the magnetic and
gravity models (Figures 9 and 10), a single body representing a
batholith can explain the observed magnetic anomalies within
the belts (KMB, EMB, and JMB, Figure 4).
Kitakami plutons. Many of the exposed Kitakami plu-
tons line up along NW trending strike-slip faults (Figure 3)
formed in response to fast, oblique subduction during the
Cretaceous [Otsuki and Ehiro, 1978]. The plutons are not offset
by the faults, but protoclastic textures observed in some of
them may have formed during faulting of semi-molten magma
(S. Kanisawa, personal communication, 1991). The textures of
the plutons and alignment with faults (Figure 3) suggest that the
faults may have helped to localize them.
Most exposed plutons are disconnected from the main
batholith. These plutons west of the batholith boundary have
been uplifted and eroded. Cretaceous volcanic rocks are still
present west of the boundary, indicating that the erosion
level is not deep. Sphalerite geobarometric data demonstrate
that the skarn deposits at the top of some of the exposed
Kitakami plutons formed at pressures between 0.9 and 1.2
kbar, that is, at depths less than 2 km [Shimizu and Shi-
mazaki, 1981]. Erosion therefore cannot account for the
difference in pluton thickness across the boundary. The
difference must be original. Some of these plutons may
represent satellite vents that tapped into the magma supply
of the main arc. The magmatic arc had shifted westward by
the Late Cretaceous [Kimura, this issue]; perhaps some of
the isolated plutons, particularly in southwest Hokkaido,
formed during the westward shift.
Cretaceous Forearc and Accretionary Prism
If Kamuikotan serpentinites cause the positive anomalies
offshore, south of Hokkaido, the Cretaceous forearc basin
rocks must underlie much of the eastern part of the modern
forearc. If they do not, there is no clear evidence for the
Early Cretaceous forearc basin and accretionary prism un-
derlying the modern forearc basin. The gravity and magnetic
models delineate nonmagnetic, high-density rocks whose
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22,180 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
depths suggest a Cretaceous sedimentary origin east of the
Kitakami plutons [Finn, 1994]. Velocities calculated from
seismic refraction data offshore [Suyehiro and Nishizawa,
this issue] are similar to the 3.5-5.8 km values found for the
Sorachi and Yezo Groups in Hokkaido [Fujii and Moriya,
1983; Ogawa et al., this issue] suggesting that these rocks
could lie under the eastern part of the forearc. The DSDP
drill hole (9, Figure 8) encountered Upper Cretaceous mud-
stone at a depth of 1145 m that may correlate with the upper
Yezo Group [Taira et al., 1983]. The mudstone was depos-
ited in a middle slope environment; the only other similar
rocks in northeast Japan belong to the Yezo group (A. Taira,
personal communication, 1994). Therefore the Upper Creta-
ceous mudstone may correlate with the upper Yezo Group
[Taira et al., 1983]. Seismic reflections from the upper
Cretaceous mudstone are found throughout the eastern half
of the forearc Iron Huene et al., 1982].
Conclusions
A New Geologic Basement Map
The minimum lateral extent of the Kitakami batholith and
plutons, and Kamuikotan serpentinites as well as qualitative
information about the shape of their buried tops can be
inferred from the geologic mapping, geophysical data, and
the models (Figure 11). The composition of the basement
geology inferred from the geophysical data requires a re-
drawing of the province boundaries in northeast Japan
(compare Figure 11 to Figure 3).
Several buried Kitakami plutons onshore were delineated
by the geophysical data (Figure 11). Others may exist in
northwest Honshu but are not detectable with the aeromag-
netic data because (1) their magnetic signature is obscured
by the overlying highly magnetic Tertiary and Quaternary
volcanic rocks, (2) they were altered and demagnetized by
the later volcanism, or (3) a combination of 1 and 2.
The new basement geologic map (Figure 11) defines much
of the composition of the western half of the modern forearc
basement as the Kitakami batholith and not the Rebun-
Kabato volcanic belt (Figure 3). The western boundary of
the batholith is on or near the Honshu coast and in central
Hokkaido. Onshore, with the exception of the strip exposed
between the Hatagawa and Futaba fracture zones (Figure 3),
exposed plutons do not form a batholith.
The location of the eastern boundary of the batholith north
of 38øN latitude is not clear. The EMB could define Kitakami
plutons (Figure 11a) connected to the main belt or com-
pletely separated from the main belt by sediments. Alterna-
tively, serpentinites may cause the observed positive anom-
alies in the EMB rather than plutons (Figure 1 lb).
The southern boundary is not clearly shown in the mag-
netic data. Positive anomalies associated with the mafic and
ultramafic rocks on land (A1, Figure 7) may belong to the
Abukuma Province [Ogawa et al., 1979] or Kitakami Prov-
ince. If the Hatagawa fracture zone (Figures 3 and 7) is
traced south to the offshore region, no clear break in the
positive anomalies can be observed (Plate 1 and Figure 7). A
zone of low-amplitude positive magnetic anomalies sepa-
rates the high-amplitude anomalies associated with mafic
and ultramafic rocks (A1, Figure 7) from positive anomalies
with wavelengths and amplitudes similar to those clearly
associated with Kitakami plutons (Plate 1 and Figures 4, 5, 7,
and 8). The southern boundary of the Kitakami Province could lie within this zone (Figure 11). Magnetic anomaly A1
is also part of the broad offshore magnetic high extending
south of Mito (Figure 1 and Plate 1). If the sources of
anomaly A1 (Figure 7) belong to the Kitakami Province, the
source of the broad high could be part of the Kitakami
batholith. If so, the Kitakami Province may extend as far
south as Mito. Alternative southern boundaries for the
Kitakami Province are shown in Figure 11. The dashed line
represents the minimum southern extent of the Kitakami
batholith, the contour the maximum extent.
In Hokkaido the KMB narrows between drill holes 4 and
7 (Figure 8) and curves west north of the Kabato Mountains
(Figure 8). The pinched-out section of the KMB coincides
with east verging thrust faults east of the Kabato Mountains
(KM, Figure 8) [Miyasaka and Matsui, 1986] and west
verging thrust faults near the Sorachi drill hole (6, Figure 8).
Collision events may have depressed the plutonic source of
the KMB below detection by the magnetometer.
The northern boundary of the Kitakami batholith is uncer-
tain. The aeromagnetic data are not extensive enough to
determine the existence of Kitakami plutons beneath Rebun
Island (Figure 1) and to the north. $egawa and Oshima
[1975] suggested that the buried Mesozoic volcanic-plutonic
source of the KMB continued to Sikhote Alin (Figure 1) on
the basis of correlation of positive aeromagnetic anomaly
belts. This proposed northern extension of the Kitakami arc
in Sikhote Alin (Figure 1) is paired with the northern
extension in Sakhalin (Figure 1) of the Cretaceous forearc
basin and accretionary prism exposed in Hokkaido [Kimura,
this issue]. However, the positive magnetic anomalies in
Sikhote Alin are associated with Paleogene volcanic rocks
and a Late Cretaceous volcano-plutonic belt [Parfenov,
1978; Takahashi et al., 1980] correlated with plutons in
southwestern Japan [Ishihara and Satoh, 1991], not Early
Cretaceous rocks. The volcanic and plutonic rocks exposed
in Sikhote Alin are younger products of the same subduction
system that caused the Kitakami plutons [Kimura, this issue]
and are not extensions of it.
I interpret that the volcanic rocks exposed in the Kabato
Mountains and on Rebun and Moneron (not shown) Islands
may be pans of the Kitakami arc not undedain by the batholith.
As mentioned earlier, exposed sections of middle Yezo sedi-
ments are undedain by silicic volcanic rocks correlated with
the Harachiyama and Kabato volcanic rocks [Girard et al.,
1991]. The volcanic rocks found in the Sorachi (6, Figure 8) and
Rumoi (7, Figure 8) drill holes are also overlain by middle Yezo
forearc sediments. The volcanic rocks may have flowed into
the forearc from the arc immediately to the west and therefore
would not belong to the Sorachi Group.
The geologic basement map (Figure 11) looks similar to
one that could be derived from the distribution of positive
aeromagnetic anomalies in California [Blake et al., 1978].
The main Cretaceous arc in California is represented by the
Sierra Nevada batholith [e.g., Hamilton, 1969]. Some of the
plutons that compose the Sierra Nevada batholith belong to
the magnetite series [e.g., Bateman et al., 1991 and refer-
Figure 11. (opposite) (a) Alternate hypotheses for the max-
imum lateral extent of the Kitakami batholith and plutons
and (b) Kamuikotan serpentines inferred from the aeromag-
netic data (Plate 1 and Figures 5, 7, and 8). Revised province
boundaries are also shown.
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FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN 22,181
45'
44 ß
43'
42 ß
41'
40'
39'
7 e
35' 100
140' 141'
I !
Kitakami province 2OO KILOMETER$
I
142' 143' 144' 141' 142'
! 140'
b
Kitakami province EXPLANATION
.................... Exposed Kitakami plutons Kitakami batholith
and buried plutons
,."-'. Kamuikotan serpentinite
21562202b, 1994, B11, Downloaded from https://agupubs.onlinelibrary.wiley.com/doi/10.1029/94JB00855 by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
22,182 FINN: AEROMAGNETIC EVIDENCE FOR A BURIED ARC, JAPAN
ences therein] and cause positive magnetic anomalies [Blake
et al., 1978] with anomalies of similar amplitude and wave-
length to those over the Kitakami batholith and plutons.
Over the Cretaceous forearc in the Great Valley of California
are positive magnetic anomalies with amplitudes up to 1000
nT [Cady, 1975; Blake et al., 1978]. These anomalies are
interpreted to be caused by serpentinized crust and mantle of
a tectonically emplaced fragment of oceanic crust [Cady,
1975]. The magnetic anomalies over these rocks have similar
amplitudes and wavelengths [Blake et al., 1978] as those
over the Kamuikotan metamorphic belt and are in a similar
position relative to the batholith. However, the tectonic
events that led to the emplacement of the Great Valley
ophiolite [Cady, 1975] are different from that proposed for
the Kamuikotan rocks [Kimura, this issue].
Tectonic Implications of the Aeromagnetic
and New Geologic Basement Maps
The many mapped strike-slip faults (Figure 3) with Creta-
ceous movement reflect the highly oblique, fast subduction
of the Izanagi plate [Engebretson et al., 1985; Taira et al.,
1983]. Features in the magnetic map evidence similar faults
within the buried Kitakami batholith. Very sharp, down-to-
the west magnetic gradients mark the western boundary of
the KMB (Figure 4) indicating a steeply dipping pluton
border. The gradients coincide with the increase in thickness
from 3 to 15 km of the Kitakami plutons [Ogawa, 1992]
(Figure 6). The MT data [Ogawa, 1992] permit the western
edge of the postulated batholith (under the steep magnetic
gradients) to dip between 60 ø and 90 ø west; it cannot slope
eastward. The straightness, steep dip, and 550-km length of
the batholith boundary suggested to Finn et al. [1992] that a
fault might have helped to localize the batholith in the Early
Cretaceous, although no fault is exposed. The exposed
plutons on the boundary are concentrically zoned with no
disruptions. The pre-Cretaceous basement is the same on
both sides of the location of the boundary. Without further
evidence for a fault, all that can be said is that the boundary
is a steep intrusive contact.
In the Late Cretaceous, the Taro (TF, Figure 5) and other
faults cut the Kitakami batholith. Sedimentary basins formed
as a result of the faulting [Minoura and Yamauchi, 1989] and
are associated with small aeromagnetic lows superimposed on
the KMB. A north-northwest trending, 100-km-long, 5- to
8-km-wide magnetic low bisects the KMB (Basin, Figure 5).
The linearity and narrowness of the low suggest that its source
is a basin controlled by a strike-slip fault, possibly northern
extensions of the fault (white line marked KF, Figure 5) that
caused the Kuji Basin (KB, Figure 5) and the Taro fault (black
and white lines marked TFE, Figure 5).
Offshore from Kinkasan (Figure 1), another positive
anomaly is bisected by a linear low (Basin, Figure 5) that
might be due to another fault-produced basin. Negative
magnetic anomalies with only a few discrete highs are
associated with an area over Sendai Bay (Figure 1) disrupted
by about 20 km of dextral movement during the Miocene
opening of the Japan Sea along a strike-slip fault that bisects
all of Honshu [Otsuki and Ehiro, 1978]. Kitakami plutons,
the sources of the small positive anomalies east of anomaly
P5 (Figure 7), may underlie the area associated with the
magnetic low but were dropped by the fault below the
detection level of the magnetometer (K. Nakamura, written
communication, 1991). The distance to the Japan Trench from the proposed eastern
edge of the Kitakami arc is between 100-140 km in the north
and 150 km in the south (Figure 11), much closer than the
average 300-400 km continental arc-trench distance [Jarrard,
1986]. The observation that the Kitakami batholith is close to
the modem trench indicates that much of the Early Cretaceous
margin, including most of its forearc, is missing.
How the material was removed is unknown; it could have
been strike-slip faulted, eroded by subduction-related pro-
cesses, or both. Oblique subduction and related sinistral
strike-slip faults of Late Cretaceous and Paleogene age
within the forearc are well-documented in central Hokkaido
[Watanabe and Kimura, 1987] and Sakhalin [Kimura et al.,
1992] and could have removed material.
Erosion of the front and base of the modern arc is
suggested by the small size of the modern accretionary prism
and subsidence of the modern forearc during the last 20 m.y.
[yon Huene et al., 1982; yon Huene and Lallemand, 1990].
The subsidence would partially explain why most of the
postulated Kitakami batholith is buried. In Hokkaido, the
Cretaceous and Paleogene accretionary complexes are
clearly truncated by the Kuril trench [Kimura, this issue].
The formation of the Japan Trench in the Paleogene may
have similarly truncated the forearc off Honshu.
Acknowledgments. Most of the work for this paper was done
while I was working at the Geological Survey of Japan in 1990 and
1991. I thank Yasukuni Okubo for hosting me and providing the
4500-foot aeromagnetic data and Masahiko Makino and Shigeo
Okuma for reading many of the data tapes for all of the aeromagnetic
data. Yasuo Ogawa furnished unpublished MT data. Yasuaki Mu-
rata helped me immensely with various computer tasks. This work
was supported by the Japanese Science and Technology Agency and
Agency of Industrial Science and Technology Fellowships and the
U.S. Geological Survey's Geothermal Program. Reviews by John
Cady, Shigeo Okuma, Rick Blakely, Randy Keller, and Gaku
Kimura greatly improved the manuscript.
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Journal of Asian Earth Sciences 184 (2019) 103968 Contents lists available at ScienceDirect Journal of Asian Earth Sciences ELSEVIER journal homepage: www.elsevier.com/locate/jseaes Trace elemental and Sr-Nd-Hf isotopic compositions, and U-Pb ages for the Cretaceous TRT triple junction offshore Japan Soichi Osozawa,b,*, Tadashi Usuki, Masako Usukid, John Wakabayashi, Bor-ming Jahn.1 Department of Earth Sciences,Graduate School of Science,Tohoku University,Sendai 980-8578Japan bKawaOso Molecular Bio-Geology Institute, Sendai 982-0807,Japan Deparnfienetnlawaniversitsveltadtiap dInstitute of Earth Sciences, Academia Sinica,128Academia Road, Section 2,Nankang, Taipei 11529,Taiwan eDepartment of Earth andEnvironmental Sciences,California State University-Fresno,Fresno,CA 93740,USA ARTICLEINFO ABSTRACT Keywords: We conducted major, trace element, Sr-Nd-Hf isotopic, and U-Pb geochronologic analyses of early Cretaceous Calc-alkaline Kitakami granitic plutons, northeast Japan. We suggest that these plutons include rocks of adakitic affinity, Zircon which indicate partial melting of an eclogitic slab. The Kitakami adakites were mostly derived from juvenile Eclogitic slab melting oceanic crustal sources, but include inherited zircons with Archean to Neoproterozoic Hf model ages. Rather Ridge subduction than being directly derived from crustal rocks of these ages, zircons may have been detrital in the voluminous Exhumation trench fill sandstone within the Jurassic northern Kitakami accretionary prism. The exhumation of older base- Transform fault ment that served as the detrital zircon source may have been triggered by ridge subduction. Although the plutons show a general southward younging from northern Kitakami (120-130 Ma), southern Kitakami (115-125 Ma), and Abukuma (100-115 Ma), restoration of post-crystallization left-slip faults inverts this spatial pattern and results in a northward-younging patern. This suggests northward migration of a trench-ridge-transform triple Kitakami adakite in time and space if the transform duplexing is restored. 1. Introduction 1999). Modern adakite is restricted to subduction settings with unusually Adakitic magma was originally proposed to have been derived by high heat-flow such as reaches associated with ridge subduction, and is slab melting (Defant and Drummond, 1990). Adakite was first reported rare compared to ordinary subduction-related calc-alkaline magmatism. from Adak Island in the Aleutians (Kay, 1978), but has subsequently A similarly high heat flow is thought to have typified the Archaean been reported from other areas associated mid-ocean-ridge subduction, Earth, resulting in subduction of hotter oceanic plates than present, slab tears, or areas with an inferred record of such subduction in the with the generation of TTG suites being a possible result (Martin et al., past. Examples include Panama along the Middle America Trench 2005). (Defant et al., 1992), Baja California (Aguillon-Robles et al., 2001), the Adakite (and TTG) are felsic and sodic igneous rocks geochemically Solomon Islands and Woodlark Basin (Konig et al., 2007), the Lau- characterized by high Sr and La compositions and relatively low Y and Tonga supra-subduction zone (Falloon et al., 2008), Daisen volcano in Yb compositions (=high Sr/Y and La/Yb; Defant et al., 1992). Gen- southwest Japan (Morris, 1995; Tokunaga et al., 2010; Zellmer et al., eration of adakitic magmas have been modeled as the product of partial 2012), and Philippines examples compiled by Castillo (2012). Archaean melting of an eclogitic slab which lacks plagioclase (Tsuchiya et al. sodic rocks of tonalite, trondhjemite, and granodiorite composition 2007; Castillo, 2012). Accordingly, coeval adakite and eclogite may (TTG) are also thought to have derived from partially melted eclogite of both be present in the same paleosubduction zone setting if they were the subducting slab (Jahn et al., 1981), and the Austral Volcanic Zone emplaced together. in Chile may represent a modern analogue of such TTG suites (Martin, In the Paleozoic ultra-high pressure metamorphic belt of western * Corresponding author at: Department of Earth Sciences, Graduate School of Science, Tohoku University, Sendai 980-8578, Japan. E-mail address: kawaoso@icloud.com (S. Osozawa). 1 Deceased 1 December, 2016. https://doi.org/10.1016/j.jseaes.2019.103968 Available online 17 August 2019 1367-9120/ @ 2019 Elsevier Ltd. All rights reserved. S.Osozawa, et al. Jourmal of Asian Earth Sciences 184 (2019)103968 China, Zhang et al. (2015) showed that source eclogite was partial initially developed as a normal fault related to the exhumation of the melted by the exhumation-related decompression. This resulted in a Sambagawa high P/T metamorphic rocks (e.g. Wallis, 1998; Fukunari rock suite that includes a melt phase (adakitic leucosomes with positive and Wallis, 2007; Kubota and Takeshita, 2008). Eu anomaly), cumulate (garnetite), and residue (amphibolitized eclo- The Kitakami zone can be divided into two subzones. The southern gite) in a single outcrop. Thus, adakite is genetically related to eclogite Kitakami zone consists of Silurian to early Cretaceous fore-arc basin strata, and pre-Silurian high P/T metamorphic, plutonic, and ophiolitic tings. basement. The pre-Silurian Hikami granite forms plutonic basement in Adakite was also reported from the Tibetan plateau associated with the southern Kitakami zone (Fig. 1), and is unconformably overlain by Indian-Asian collision (Chung et al., 2009). Cretaceous granitoids of the Silurian arkosic sandstone (Murata et al., 1974). The pre-Permian strata Yangtze River Belt were attributed to partial melting of thickened/de- contain pyroclastic rocks derived from a neighboring volcanic arc, and laminated lower continental crust (Yan et al, 2015), but recent work has the Triassic strata locally intercalates silicic tuff. The Upper Jurassic reinterpreted these plutons as a product of slab melting of subducted shallow marine mudstone and sandstone are overlain by the Early oceanic plate (eclogite) that generated adakitic melts (Deng et al., Cretaceous terrestrial andesitic and dacitic pyroclastic rocks correlated 2016). with the Harachiyama Formation of the northern Kitakami zone men- Cretaceous adakitic rocks, including major granitic plutons are in- tioned below. truded into the rocks of the Kitakami zone, northeast Japan (Tsucjhiya The northern Kitakami zone primarily consists of a Jurassic accre- and Kanisawa, 1994; Tsuchiya et al., 2007, 2015). These Early Cre- tionary prism and accordingly lacks a continental basement. The taceous granitoid intrusions were spatially and temporally associated Hayachine fault (Fig. 1) separates the northern and southern Kitakami with an inferred subduction zone, and they intrude a Late Jurassic ac- zone, and it may have been reactivated as a sinistral fault. The Lower cretionary prism. Thus, the Kitakami adakitic plutons are more likely to Cretaceous Harachiyama Formation consists of altered andesite and have been derived from a subducted eclogitic slab than from partial dacite closely associated with granitic plutons (Kanisawa, 1974) that melting of the lower continental crust. have been shown to be adakitic (Tsuchiya and Kanisawa, 1994). Close In an earlier study of the Kitakami zone, we analyzed the de- to the sample locality of the Harachiyama Formation (Fig. 1), terrestrial formation and exhumation of the granitic plutons and proposed that andesitic pyroclastic rocks and lavas of the Harachiyama Formation they were generated and emplaced as a result of a mid-ocean-ridge overlie Lower Cretaceous marine mudstone of the Omoto Formation, a subduction event (Osozawa et al., 2012). In this study, we present small-scale fore-arc basin formed over the Late Jurassic accretionary major and trace element data, Sr-Nd isotopic data, and U-Pb age data prism. Accordingly, based on above regional field relationships, the combined with zircon Hf isotopic data for the Kitakami plutons. U-Pb Harachiyama adakitic rocks were also formed in a subduction zone age data for the Kitakami plutons was recently published (Tsuchiya setting. et al., 2015), and we integrate a review of these ages with our own data The Kitakami granitoids in the northern Kitakami zone have been in our evaluation of the genesis and emplacement of these plutons. classified as adakitic granitoid, calc-alkaline granitoid, and gabbroid- Similarly, Hirahara et al. (2015) reported whole rock Hf isotope data granitoid from east to west (Fig. 1; Tsucjhiya and Kanisawa, 1994; from two samples of the Kitakami plutons, so these data are also useful Tsuchiya et al., 2007, 2015). The Tono, Senmaya, and Hitokabe plutons to consider in conjunction with our own data. We also analyzed samples in the southern Kitakami zone (Fig. 1) are giant zoned plutons, and from the Abukuma and Niigata plutons in the Tohoku District of Japan, their central parts are silicic and adakitic (Mikoshiba and Kanisawa, and the isotopic study by Hirahara et al. (2015) was focused on these 2008). These plutons in the southern Kitakami zone are classified as plutons. We evaluate the data on these plutons in the context of re- adakitic, including the foliated Kesengawa body (Osozawa et al., 2012), gional structural and geochronological data including that from the but there are also many gabbroid-granitoid bodies (Fig. 1; Tsuchiya Sambagawa eclogite， southwest Japan (e.g, Osozawa and et al., 2007, 2015). Wakabayashi, 2017), and propose a tectonic model for the genesis of -d n s n p s s the adakitic rocks that includes mid-ocean-ridge subduction and plied to the Abukuma and Niigata plutons (Fig. 2); these are described transform faulting during early Cretaceous time (Osozawa, 1997b). by granitoid names below (Table Si). The Nishitagawa body, one of the Nigata plutons, is structurally included in the Asahi zone, and the 2. Geological setting Iwafune body is in the Ashio zone (Hirahara et al., 2015), and these two local zones are extensions of the Sanyo zone of southwest Japan Cretaceous granitoid plutons are components of both the Sanyo- (Osozawa, 1997a). Ryoke zone of southwest Japan and the Kitakami-Abukuma zone, northeast Japan (Fig. 1, inset). Based on K-Ar hornblende and biotite cooling ages of these plutons, evaluation of regional geology and global 3. Samples and methods plate models, Osozawa (1997a) proposed that they were exhumed following two episodes of ridge subduction in the early and late Cre- Sample localities are shown in Figs. 1 and 2. 49 samples (32 also for taceous, respectively (Osozawa, 1997a). Early Cretaceous structures U-Pb and Hf analyses) were analyzed for the Kitakami zone, 19 samples such as folds and slaty cleavages related to exhumation of the Kitakami (6 also for U-Pb and Hf analyses) for the Abukuma zone, and 7 (no plutons were recently described (Osozawa et al., 2012). sample for U-Pb and Hf analyses) for the Nigata plutons. In Table S1, The present distribution of geologic units in southwest and north- sample numbers, rocks or powders, latitudes and longitudes, granitoid east Japan has been significantly affected by sinistral transform faulting classification names, pluton names, granitoid names by CIPW norma- that acted between the two ridge subduction events described above. As tive Q-ANOR classification below, collectors, XRF for major elements, a consequence, Cretaceous and older geologic units of northeast Japan ICP-MS for trace elements, TIMS for Sr-Nd isotopes, CL images of zir- originally lay south of the rocks in present southwest Japan (Osozawa, cons, LA-ICP-MS for U-Pb ages, LA-ICP-MS for zircon Hf isotopes, and 1994, 1997a, 1997b). The Tanagura sinistral fault, once a primary machine analysts are shown. For methodological details, see Jahn et al. transform fault, bounds southwest and northeast Japan. The Hizume- (2014) and Appendix (Method Details). Kesennuma and Futaba faults are branches of this fault (Figs. 1 and 2; Osozawa, 1997b). Southwest Japan is divided into the inner and outer zones bounded by the Kurosegawa tectonic zone, major sinistral strike- slip fault zone that links to the Tanagura fault. The Median Tectonic Line was not a primary transform fault at this time, and it may have 2 S. Osozawa, et al. Journal of AsianEarth Sciences184(2019)103968 dehydration of a subducting slab and coupled hydration of the overlying KTKM-11 Hashi- Okawame Mine Oguni gami KT708 Northern BJ-13-101 Kuki quakes and chemical element recycling (e.g., Hacker et al., 2003; 142°E KTKM-08 Kitakami zone Japan. Nozaka (2005) has shown that mylonitic shear zones within the Kosode KTKM-24 BJ-13-102 Sakainokamidake Otanabe 40°N KTKM-09 KTKM-02 Tenjinmor Tanohata Himekami ano and later formation of the schistose serpentinites. KTKM-06 0024-4937/$ - see front matter 2013 Elsevier B.V. Allrights reserved. occurred under a wide range of temperature conditions. Taro Nedamo KTKM-03 subbelt KTKM-12 Yayakomlor Omoe Hayachineophiolite KTKM-13 Miyako KTKM-14 Hizume-Kesennuma fault Tono KTKM-15 70K103 Kurihashi KTKM-16 n 70K45 Kushi KTKM-17 Tono KT653 cea KTKM-04 KT654 70K568 KT640 0 KT627C Hitokabe Kadokamiwa C KTKM-23 KT632 Akagane Mine achinefaultGoyosan 71K220 KTKM-18 KTKM-27 a Motai Ryori" KTKM-28 P metamorphic Sen- 39°N rocks maya Hikami granite Hondera KTKM-19 KTKM-05 Department of Earth Sciences Okayama University, Okayama 700-8530,Japan Tabashine BJ-13-106 731103 Senmaya 70K157 Hirota Southern Orikabe KTKM-22 KTKM-25 Kitakami Kesengawa KTKM-26 KTKM-21 70K193 zone BJ-13-107 70K156 70K154 25 km N An important example of such linkage is the Happo ultramafic OYA12 OYA13 Oshika Kinkasan Legend North Kitakami Belt Jurassic complex entire sequence of alteration and tectonic processes in the supra- Kitakami Pluronic rocks of Early Cretaceous Nedamo Belt Niigata zone geologic proceses such as arc magmatism, dehydration-induced earth- Adakitic granitoid NE Japan complex Calc-alkaline South Kitakami Belt granitoid 2004; Scambelluri et al., 1995), the serpentinites seem to have been Cretaceous Gabbroid -granitoid Jurassic zone Triassic Volcanicrocks Permian Early Cretaceous pre-Permian posed exposures of peridotite mylonites and schistose serpentinites complex Fig. 1. Locality map of the Kitakami plutons. Pluton names and specimen numbers are shown. S.Osozawa, et al. Jourmal of Asian Earth Sciences 184 (2019) 103968 Niigata SWJapan 71Ni2 149°30' N 71Ni4 Nishitagawa lwafune ctor 71Ni17 71Ni16- Abukumazone 71Ni11 NE Japan 38° 71Ni21, 25 km 71Ni24 Late Cretaceous granitoids a acific Sheared granitoids 113 Ocean Neogene granitoids 115 A15,A38 101 113 A41 A34 118 109 105 104 100 A36 107 A35 111 A52 111 112108 A48,A49 A1, A4 . 103 106 A53 Ishikawa 109】 119Ab-5 121 10410.17.02 107 10.17.03 37° Ab-4( 110 Samegawa Ab-3 108 Ab-2 108 25 km Ab-1 Cretaceous granitoids Gosaisho-Takanuki metamorphic rocks Miyamoto Fig. 2. Locality and age map of the Abukuma and Nigata plutons. The Nigata plutons are corresponding to the Sanyo plutons, southwest Japan and its northeastern extension. Pluton names and specimen numbers are shown. Oval age: this study, orange: Kon and Takagi (2012) and Kon et al. (2015), blue: Takahashi et al. (2016), green: Ishihara and Orihashi (2015). The based geological map is after Kano et al. (2002) and Ishihara et al. (1998). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.) 4.Result and other secondary minerals, reflecting hydrothermal alteration, so major elements should be regarded with caution for chemical dis- 4.1. Discrimination by major elements crimination among these samples. The other granitoid samples that we analyzed are fresh and lack visible evidence of alteration, so chemical 4.1.1. Major element compositions are shown in Table S2 modification of such samples is probably minor. The following dia- The Hikami granite and the Harachiyama volcanics contain chlorite grams are common to those of Fiannacca et al. (2015). On the Al2O3 vs S.Osozawa,et al Jourmal of Asian Earth Sciences 184 (2019) 103968 4.0 1.0 Kitakami ● Calc-alkaline granitoid Adakiticgranitoid 0.9 3.5 A-type ●Gabbroid-grantoid + Hikami granite Abukuma Mg 0.8 Ferroan 3.0 ▲Niigata + 0 0.7 Fe* Metaluminous Peraluminous 2.5 S-type Z 0.6 A 2.0 FeO 0.5 1.5 0.4 I-type 0.3 1.0 45 50 55 60 65 70 75 Peralkaline 80 85 I-type S-type 15 0.5 A-type 0.5 0.7 0.9 1.1 1.3 1.5 10 A/CNK CaO S-type Fig. 3. Al2O3 vs (CO2 + Na2O + K2O) diagram. 1 5 K0 (CO2 + Na2O + K2O) diagram (Fig. 3), granitic rocks are classified into Y + I type granite, except for the Hikami S type granite and including the O Kitakami Silurian sandstone. Excluding the Hikami granite, rocks plot in the Na2 Calc-alkaline granitoid Adakiticgranitoid 5 Gabbroid-granitoid The CIPW normative Q'-ANOR (quartz, anorthite, oligoclase, albite) Calc + Hikami granite Calcic I-type classification diagram (Fig. 4) shows a wide diversity of rock types, Abukuma ▲Nigata which range from gabbro to syenogranite, and the compositional trends -10 65 45 50 55 60 75 80 85 range from calc-alkalic to calcic. On the FeOtotal/(FeOtotal + MgO) vs SiO2 diagram (Fig. 5) and SiO, (%) Na2O + K20 - CaO vs. SiO2 diagram (Fig. 6; Wilson, 1989), the gran- Fig.5. Top: FeOtotal/(FeOtotal + MgO)vsSiO2 diagram, itoids plot as magnesian and calc-alkalic to calcic with some exceptions, bottom: Na2O + K20-CaO vs SiO2 diagram (Frost et al., 2001). and in I type granite field. On the K2O vs SiO2 diagram (Fig. 7; Rickwood, 1989), the granitoids define high-K or medium-K calc-al- kaline trends, with the exceptions of four analyses of the Orikabe and 50 Granodiorite Kitakami Monzogranite Calc-alkaline granitoid ● Gabbroid-granitoi An) ● Adakitic granitoid Gabbroic rocks + Hikami granite Tonalite 40 Abukuma + Calcic Granitic rocks b Gabbroic rocks + 6 30 + Q AIkal 20 Q × 0 10 1 Alkali-feldspa Quartz Quart quartz syenite Alkali-feldspar Syenite 0 0 20 40 60 80 100 ANOR = 100 x An / (Or + An) Fig. 4. CIPW normative Q'-ANOR (Q: quartz, An: anorthite, oligoclase Or: orthoclase, Ab: albite) classification diagram (Streckeisen and Le Maitre, 1979). The boundaries used are those more recently proposed by Whalen and Frost (2013). S. Osozawa, et al. JournalofAsianEarthSciences184(2019)103968 Ultrabasic Basic Intermediate Acid Kitakami 16 ● Calc-alkaline granitoid ●Adakitic granitoid Gabbroid-granitoid 14 Nepheline Gabbroicrocks + Hikami granite syenite Abukuma Grar 12 Gabbroic rocks % Nigata 三 Syenite 10 .+.. Alkaligranite +K2O Alkalic Syenodiorite Subalkalic Na2O ljolite 6 Gabbro- Granite Quartz Diorite diorite Gabbro (Granodiorite) 0 40 50 60 70 80 SiO2 (wt %) Fig. 6. (Na2O + K2O) - CaO vs. SiO2 diagram (Cox et al.,1979 adapted by Wilson, 1989 to plutonic rocks) Himekami bodies that plot in the Shoshonite series. zone, and the Tono, Hitokabe, and Senmaya bodies in the southern Kitakami zone are classified as adakitic plutons. Chondrite-normalized REE patterns of adakites typically show 4.2.Discrimination by trace elements broad concave-up profiles, with very low heavy REE (HREE), relatively high LREE contents, and a positive Eu anomaly (Martin, 1999; Falloon 4.2.1. Trace element compositions are shown in Table S2 et al., 2008; Tokunaga et al., 2010). This pattern is found in two to- On the Sr/Y vs Y diagram (Fig. 8; Defant and Drummond, 1990; nalites from the Miyamoto and Ishikawa bodies of the Abukuma plu- Martin et al., 2005; Castillo, 2012; Tsuchiya et al., 2015), granitoids are tons, and tonalite and granodiorite from the Nishitagawa body of the plotted in adakite or island arc andesite- dacite- rhyolite field. In Fig. 1, Nigata plutons (Fig. 9). Although we found only higher HREE patterns which combines our data with data from Tsuchiya et al. (2007), the with negative Eu anomalies (Cole and Stewart, 2009) from the calc- Hashigami, Tanohata, and Miyako bodies in the northern Kitakami 6 Kitakami ● Calc-alkaline granitoid ●Adakitic granitoid Gabbroid-granitoid 5 Gabbroicrocks + Hikami granite Abukuma Granitic rocks Gabbroic rocks 4 ▲ Nigata S 0 1 + Low-K(tholeite)series 0 40 50 60 70 80 90 SiO. (wt%) Fig. 7. K2O vs SiO2 diagram. The boundaries of each series: Rickwood (1989). chromian compositions in equilibrium with primary olivine and pyrox- JournalofAsianEarthSciences184(2019)103968 300 Kitakami; Kitakami; ● Calc-alkaline granitoid Calc-alkalinegranitoid 50 Adakitic granitoid ●Adakitic granitoid amount of clinopyroxene, although the fine-grained aggregates locally ●Gabbroid-granitoid Abukuma Abukuma 250 ▲Nigata Nigata 2007;2015) 40 run off out of the pseudomorphs to form veins connecting the pseudo- zoned plutons. 200 2007;2015) Adakite 1984; Ishi et al., 1992) as do those of the Happo, Tari-Misaka, Ohsayama b 150 Y pseudomorphs after extremely elongated amphibole crystals, most of Adakite 20 100 mains are pseudomorphs after orthopyroxene with a subordinate (osnueh (Touchia1s a al (Fig. 4), and are cut by serpentine but never cut it. Olivine in relatively the Oeyama ophiolite could be masked by subsequent serpentinization. ADR ADR gesting their three-dimensional interconnection via hidden veins, and + N-MORB 0 minerals after phlogopite are optically and chemically different from ex- 20 30 phibole crystals (Fig. 3c, d). Compared with tremolitic amphibole, 0 1 2 3 4 Y (ppm) Yb (ppm) (Castillo, 2012) is shown. alkaline granites of Kitakami plutons (Fig. 9), an adakitic REE pattern 4.4. Pb ages and Hf isotopic composition from zircons with a positive Eu anomaly was reported from the central silicic part of the Tono body (Mikoshiba and Kanisawa, 2008). The calc-alkaline Cathodoluminescence images of the studied zircons do not indicate pattern with negative Eu anomalies characterizes most of the Abukuma significant alteration, except for KTKM-20 from the Hikami granite (Fig. plutons and the Iwafune body of the Niigata plutons. The calc-alkaline S12AB). Weighted mean U-Pb ages for the Kitakami and Abukuma trend of the Hikami granite and its large sample to sample composi- granitoids are shown on zircon on U-Pb concordia age diagrams (Fig. tional range may be partly an artifact of hydrothermal alteration, but S13AB). The U-Pb ages are shown in Table S4. the negative Eu anomalies are clearly primary. A gabbroid from the Our U-Pb age data supplemented by previous data of Tsuchiya et al. Tono body, Kitakami plutons, is LREE depleted with a mild positive Eu (2015) are shown in Fig. 12 (Kitakami) and Fig. 2 (Abukuma; Nigata anomaly, as shown by Mikoshiba and Kanisawa (2008). plutons are not analyzed). Plutons of the northern Kitakami zone are Primitive mantle normalized trace element abundance patterns older than 119 Ma (~130 Ma), and the adakitic, calc-alkaline, and (Fig. 10; Aguillon-Robles et al., 2001; Falloon et al., 2008; Cole and gabbroic granitoids (Fig. 1) have the same age range. Plutons of the Stewart, 2009) show lower abundances of trace elements as reflected in eastern southern Kitakami zone,east of the Hizume-Kesennuma fault very low Rb, K, and Ti abundances, and show typical subduction-re- are younger than 119 Ma (~117 Ma) in its northern part, and older than lated negative Nb and Ta anomalies relative to La. A LIL component 120 Ma in its southern part. Plutons of the western southern Kitakami shift is recognized in the altered Harachiyama volcanics, and the Hi- zone west of the Hizume-Kesennuma fault show similar spatial-tem- forearc peridotites, rather than abyssal peridotites (Dick and Bullen, poral patterns (128-112 Ma). The Tono zoned pluton in the northern part of the western southern Kitakami zone was dated 113 Ma for the central silicic and adakitic part, and 119 Ma for the marginal part 4.3.Nd-Sr isotopic composition (Tsuchiya et al., 2008; Mikoshiba and Kanisawa, 2008). Plutons of the 4.3.1. Nd-Sr isotopic compositions are shown in Table S3 these northern parts of the southern Kitakami zone (118-115 Ma), and On the eNd(t)-(87Sr/86 Sr)i diagram (Fig. 11; Cole and Stewart, plutons from the main Abukuma zone (Fig. 2) are younger than 115 Ma 2009), the granitoid data lie in the enriched extension of the mantle (~96 Ma). array. The Abukuma granitoids show enrichment relative to the de- The Hinomiko, Himekami, Sakainkamidake, and Kesengawa plutons pleted Kitakami granitoids, and the Niigata granitoids are more en- have recycled zircons of Paleozoic and Proterozoic (Fig. 13, Table S4). riched than the Abukuma granitoids, concordant to a recent study by U-Pb ages of the Hikami granite are 442-449 Ma (Fig. S13A; Table Hirahara et al. (2015). The Hikami granite samples are as enriched as S4; latest Ordovician to earliest Silurian). The oldest fossil date from the the Nigata granitoids, with the exception of two highly altered samples Silurian sedimentary cover is the Llandovery (443.4-433.4 Ma) (Murata that depart significantly from the mantle array. et al., 1974), and this supports the latest Ordovician U-Pb age of the Hikami granite. The Hikami granite contains Proterozoic recycled zir- cons (Table S4). S. Osozawa, et al. JournalofAsianEarthSciences184(2019)103968 1000 1000 Kitakami; Kitakami; Gabbroid-granitoid Adakitic granitoid 100 100 10 10 (Tsuchiya etal, 2007) Ccentralfacisc La Ce Pr NdSmEu Gd Tb Dy Ho Er TmYb Lu La Ce Pr NdSmEu Gd Tb Dy Ho Er TmYb Lu 1000 1000 Kitakami; Kitakami; Gabbroic rocks Calc-alkaline granitoid 100 100 Chondrite 10 10 La Ce Pr Nd SmEu GdTb Dy Ho Er TmYb Lu La Ce Pr Nd SmEu GdTb Dy Ho Er Tm Yb Lu Rock 1000 1000 Kitakami; Abukuma; Hikami granite Granitic rocks 100 100 10 10 0 La Ce Pr Nd SmEu GdTb Dy Ho Er TmYb Lu La Ce Pr Nd SmEu GdTb Dy Ho Er TmYb Lu 1000 1000 Niigata Abukuma; Gabbroic rocks 100 100 10 10 La Ce Pr NdSmEu GdTb Dy Ho Er TmYb Lu La Ce Pr Nd SmEu GdTb Dy Ho Er TmYb Lu Fig. 9. Chondrite normalized REE pattern. Normalizing values are from Sun and McDonough (1989). S. Osozawa, et al. Journal of Asian Earth Sciences 184 (2019) 103968 500 500 50 50 5 5 1 Kitakami; Kitakami; Gabbroid-granitoid Adakitic granitoid 0 Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu Ba UNb La Sr PSm Ti Dy YYb BaUNbLa SrPSm TiDyYYb 500 500 50 50 A mantle 5 5 Kitakami; 1 Kitakami; Gabbroic rocks Calc-alkaline granitoid Rock / Primitive 0 Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu BaUNbLaSrPSmTiDyYYb Ba UNbLa Sr PSm Ti DyYYb 500 500 50 50 5 Kitakami; Abukuma; Hikami granite Granitic rocks 0 RbThKTa Ce Nd ZrEu TbHo ErLu Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu Ba UNb La Sr PSm Ti Dy YYb BaUNbLaSrPSmTiDyYYb 1000 500 100 50 10 5 1 Niigata Abukuma; Gabbroic rocks O 0 Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu Rb Th K Ta Ce Nd Zr Eu Tb Ho Er Lu BaUNbLaSrPSmTiDyYYb BaUNbLaSrPSm TiDyYYb Fig. 10. Primitive mantle normalized trace element abundance pattern. Normalizing values are from Sun and McDonough (1989). S. Osozawa, et al. Jourmal of Asian Earth Sciences 184 (2019) 103968 10 Kitakami Cenozoicsubducted 8 Slab-related adakite Calc-alkalinegranitoid (Zhao et al., 2008) ●Adakitic granitoid ●Gabbroid-granitoid 6 Marginal facies of the Gabbroic rocks adakiticzonedplutons + Hikami granite 4 (Tsuchiya et al., 2007) Abukuma Granitic samples 2 Gabbroic rocks Nigata 0 3 -2 -4 Central facies of the adakiticzonedplutons 9- (Tsuchiya et al., 2007) -8 -10 0.702 0.704 0.706 0.708 0.710 87 Fig. 11. eNd(t)-(87Sr/86 Sr)i diagram. MORB and mantle array are after Zhao et al. (2008), and the data for Cenozoic oceanic slab-related adakites are also in Zha0 et al. (2008), taken from Defant et al. (1992), Kay et al. (1993), and Aquillon-Robles et al. (2001). Zircon Hf isotopic compositions are shown in Table S5. Hf isotope 2012). Alternatively, the Jurassic prism contained voluminous trench evolution diagram (Fig. 13) indicates that plutons of the Northern Ki- r ss i n s takami zone (Jurassic accretionary prism) tend to have recycled zircons proterozoic to Archaen detrital zircons. (Hinomiko: model ages of 4000-1500 Ma, Himekami: 2500-800 Ma, Although old recycled zircons were not found from the Kitakami Oguni: 1500-1000 Ma, Tanohata: 1200 Ma, Sakainokamidake: granitoids in the southern Kitakami zone, the inferred continental 1000-800 Ma, Otanabe: 1000 Ma, Tanohata: 1100 Ma). The Cretaceous basement there may also lack Archean and Neoproterozoic rocks. The plutons of the Southern Kitakami zone are juvenile origin, but with Hikami granite has U-Pb age of 442-449 Ma (late Ordovician), and recycled zircons that have Hf model ages of less than 1200 Ma (Ke- sengawa). In contrast, the Ordovician Hikami granite has recycled zir- other pre-Silurian unit of the southern Kitakami zone is the Hayachine cons of 3000-1300 Ma. ophiolite and adjacent high P/T metamorphic rocks, also formed at subduction zone setting without continental basement (Osozawa et al., 5. Discussion 2012). In addition, the Silurian to Carboniferous sedimentary cover consists of arc type volcaniclastics, and including the Hikami granite, 5.1. Juvenile granitoids of the Kitakami plutons? they represent volcanic arc or the frontal arc (Osozawa et al., 2012). The Archean and Neoproterozoic continental crust is inferred to have The high Sr - low Y, and low-HREE patterns with positive Eu been west of the southern Kitakami zone, so recycled zircons of the anomalies, that characterize some of the samples, reinforces previous Hikami granite may have been derived from detrital zircons from the data of Tsuchiya et al. (2007) and Mikoshiba and Kanisawa (2008), and clastic sedimentary rocks that it intruded. demonstrates the existence of adakitic granitoids in the Kitakami zone, Aoki et al. (2015) obtained Archean and Neoproterozoic U-Pb ages as well as the Abukuma zone and the Niigata granitoids. These char- from the Kurosegawa zone granites of southwest Japan (the Mitaki acteristics suggest that the adakitic magma was derived from partial granite correlated to the Hikami granite), and interpreted that the zir- melting of the juvenile oceanic crust that had been metamorphosed to cons were recycled from terrigenous sedimentary rocks, such as sand- eclogite. stone, that were deposited on continental crust containing Precambrian The other non-adakitic juvenile granitoids have higher HREE with a basement rocks in South China. Such a recycling process was also negative Eu anomaly and primitive-mantle-normalized trace element s sd a ro a o p pattern with negative Nb and Ta anomalies. Such rocks are character- west Japan (Jahn, 2010). ized as calc-alkaline granitoids, including gabbroid-granitoids, and re- Although the Nishitagawa body, Nigata plutons, are adakitic on the basis of their REE pattern (Fig. 9), they are enriched on the Nd-Sr mantle wedge above the subduction zone. Gabbroids have lower LREE diagram (Fig. 11) similar to other plutons of the Sanyo zone. Such with a positive Eu anomaly, suggesting plagioclase fractionation in enrichment relative to the mantle array (plotted above the array; cumulates. Fig. 11) was ascribed to mantle-lower crustal assimilation (Hirahara Juvenile granitoids in the northern Kitakami zone, however, contain sd e n recycled zircons of Paleozoic to Proterozoic and Archean to fected by assimilation than the Kitakami plutons (Fig. 11). Neoproterozoic Hf model ages, and this conflicts with their juvenile geochemical character. The zircons were probably not directly derived 5.2. Relationship of adakites to Sambagawa eclogite?: Infant subduction from Archean to Neoproterozoic crustal basement because the Late zone Jurassic northern Kitakami zone was relatively a young accretionary prism lacking the continental basement before emplacement of the Eclogite facies metamorphism of the Sambagawa metamorphic belt Early Cretaceous Kitakami granitoids (see Fig. 14, Osozawa et al., of southwest Japan, may have taken place during at least two 10 S. Osozawa, et al. Jourmal of Asian Earth Sciences 184 (2019) 103968 126 125 142E North 124 123 Kitakami Belt 123 123 122 125 40°N 32? 122 127 121 124 122 121 125 124 NedamoBelt 124 123 n Ocear 127 124 121123 121 acific 119 South a Kitakami Belt 118 1202 P 25 22 119 39°N 13 122 112 Legend Zircon U-Pb ages ≥ 125 Ma 120-125 Ma Japan 115-120 Ma 25km 110-115 Ma Volcanic rocks 厂 Early Cretaceous Plutonicrocks of Early Cretaceous Adakitic granitoid Calc-alkaline granitoid Gabbroid - granitoid Plutonic rocks ofpre-Permian Hikami granite 15 Fig. 12. Zircon U-Pb age map of the Kitakami plutons. The filled oval ages indicate our data. The unfilled oval age is from Tsuchiya et al. (2015), and the southernmost 115 Ma data is after Orihashi and Ishihara (2015). subduction metamorphic episodes or over a range of time (Osozawa Takasu (1989). Corroded zircons in quartz eclogite indicate that the and Wakabayashi, 2017), at 120-110 Ma (0kamoto et al., 2004) and at subsequent partial melting after the peak eclogite metamorphism oc- curred from 116 to 104 Ma (Arakawa et al., 2013); such partial melting The older ages were obtained from quartz eclogite, interpreted as a may have yielded an adakitic melt. The 120-110 Ma metamorphic age volcaniclastic rock, that contained recycled (detrital) zircons as old as (Okamoto et al., 2004) of the older Sambagawa eclogites corresponds to Proterozoic (Okamoto et al., 2004). Various metamorphic and field formative age of the Kitakami adakite, and the source of the adakite relationships of the dated eclogites were interpreted to suggest a pos- may be locally preserved in the Sambagawa belt. sible olistostrome block origin for the older eclogites (Osozawa and REE pattern of Sambagawa quartz eclogite ：was reported by Utsunomiya et al. (201l). If this eclogite partially melted, then the 11 S.Osozawa, et al. Jourmal of Asian Earth Sciences 184 (2019) 103968 20 10.1 10 1000Ma EHfCHUR 0 2000Ma 10 3 -20 3000Ma 4000Ma Kitakami Calc-alkaline granitoid - Ab-3 Abukuma 30 6.35 Gabbroid-granitoid BJ-13-101 ■KTKM-04 × KTKM-21 X KTKM-05_× KTKM-11 Gabbroic rocks O KTKM-02 KTKM-10 X KTKM-23 ●Ab-4 _KTKM-06KTKM-24 50.7 0.01 Ab-5 ▲KTKM-17 0.44 KTKM-07 Cpx -40 0 200 400 600 800 1000 1200 1400 1600 1800 2000 206pb- age (Ma) Fig. 13. Hf isotope evolution diagram. DM: Depleted mantle, CHUR: chondritic uniform rese rvoir. Dashed lines: Hf evolution lines of 1000, 2000, 3000, and 4000 Ma crust. Average Lu/Hf of crust: O.015. residue of hypothetical adakitic melt has relatively higher LREE with a north; 100-115 Ma (Abukuma), 115-125 Ma (southern Kitakami; east weak Eu positive anomaly, comparable to the REE profile of residual and west of the Hizume-Kesennuma fault is duplicated in age), and amphibolitized eclogite reported by Zhang et al. (2015) for the ultra- 120-130 Ma (northern Kitakami; east and west of the Hayachine fault is high-pressure metamorphic belt of western China. Note that the REE partly duplicated in age) (Fig. 14). This diachronous adakitic plutonism (o) a a z eo suggests that juvenile slab migrated from south to north, and this event characterized by a negative Eu anomaly, that may reflect mixing with no n pa pyroclastic or continental detrital material (Okamoto et al., 2004). Fig. 14). The 110-120 Ma age of the older Sambagawa eclogite (north In the Sambagawa zone, we suggested that the quartz eclogite may of the Kurosegawa tectonic line; therefore included in the inner zone of have been formed during subduction initiation (Osozawa and southwest Japan) may be consistent with this model. Wakabayashi, 2017). The adakitic granitoids in the Kitakami zone may The T(trench)T(trench)R triple junction commonly envisioned is have formed during such an event. instead proposed as a T(trench)RT(transform) triple junction (Fig. 14). Osozawa, 1994, 1997a, 1997b) proposed that the plate motion change of the Izanagi plate parallel to the trench generated transform fault 5.3.Ridge subduction followed by transform faulting: Migrating TRT triple plate boundary during 115 to 100 Ma. However, present U-Pb age data junction shows that this did not take place synchronously along the length of the trench, and we propose a California type event with a northward-mi- The Tanagura tectonic line is the major sinistral fault bounding grating TRT triple junction (Fig. 14) (e.g., Atwater, 1970; Furlong, northeast Japan (Kitakami and Abukuma zones) and southwest Japan 1984). We proposed the upright extrusional domes along a restraining (Nigata gnanitoids and Sanyo zone) (Fig. 2), and originally a transform bend of a forearc strike-slip fault system for the Triassic blueschist fa- fault (Osozawa, 1994, 1997b). The radiolarian ages of cherts in accre- cies subduction complex rocks in the Ishigaki-jima island, Ryukyu tionary complexes indicate that northeast Japan and southwest Japan (Osozawa and Wakabayshi, 2012), but this transcurrent movement was (subdivided into the inner and outer zones by the Kurosegawa tectonic the unrelated post Eocene event. line) were duplicated by transform faulting between ca.115 and 100 Ma At ca. 120 Ma, original position of the Kitakami and Abukuma (must have taken place before 90 Ma from maximum date of Cretaceous plutons was close to south China adakitic plutons formed by the infant -ss on d ( u sd Izanagi plate subduction (Deng et al., 2016; Fig. 14). The Kitakami and tion) configuration from south to north was the outer zone, northeast Abukuma plutons were translated northeastward more than 1oo0 km Japan, and the inner zone (Osozawa, 1994). from the original formative position of south China after 115 to 100 or Within northeast Japan, the Hatagawa and Furaba faults (Fig. 2) 90 Ma. The estimated trajectory of the Izanagi plate is permissive of this and the Hizume-Kesennuma and Hayachine faults (Fig. 1) are such si- large displacement (Osozawa, 1997b). Recycled zircon U-Pb dates of nistral faults related to the above transform system, and the pre- the Hikami granite of northeast Japan, including the Mitaki granite of transform relationship of units from south to north was the northern inner zone, southwest Japan (Aoki et al., 2015), suggest that these Kitakami zone, southern Kitakami zone east of the Hizume-Kesennuma plutons were also close to the Precambrian basement of south China at fault, southern Kitakami zone west of the Hizume-Kesennuma fault, and Ordovician time. the Abukuma zone. By using the younger (middle Cretaceous to Paleogene) cooling ages If such sinistral faulting is restored, the U-Pb ages young to the 12 S. Osozawa, et al. JournalofAsianEarthSciences184(2019)103968 >100 Ma adakite Abukuma north China includingHikamigranite Farallonplate alkaline south China adakite fromslab adakite south China >110 Ma slab subducting eclogte) slab window Farallon plate 、trench-ridge-transform such as triple junction Tanagura let fault adakite south China lateral >120 Ma strike-slip_ Izanagi plate Fig. 14. Tectonic model of proto Japan at 100-120 Ma. of the Sanyo granitic plutons (Osozawa, 1994, 1997a) and Sambagawa References metamorphic rocks (Osozawa and Pavlis, 2007; Osozawa and Wakabayashi, 2015), we showed the diachronous exhumation related Aguillon-Robles, A., Caimus, T., Bellon, H., Maury, R.C., Cotton, J., Bourgois, J., Michaud, to the nothward migration of the TTR triple junction (Kula-North New F., 2001. Late Miocene adakite and Nb- enriched basalts from Vizcaino Peninsula, Mexico: indicators of East Pacific Rise subduction below southern Baja California. Guinea plates; Osozawa, 1994, 1997a). The Nigata plutons may have Geology 29, 531-534. been related to this younger event of ridge subduction, and the younger Aoki, K., Isozaki, Y., Yamamoto, A., Sakata, S., Hirata, T., 2015. Mid-Paleozoic arc dates of the Sambagawa eclogitic metamorphism (89-88 Ma) may be granitoids in SW Japan with Neoproterozoic xenocrysts from South China: New zircon U-Pb ages by LA-ICPMS. J. Asian Earth Sci. 97, 125-135. related. Arakawa, M., Okamoto, K., Yi, K., Terabayashi, M., Tsutsumi, Y., 2013. SHRIMP U-Pb dating of zircons related to partial melting in a deep subduction zone: Case study from the Sanbagawa quartz-bearing eclogite. Isl. Arc 22, 74-88. Acknowledgements Atwater, T., 1970. Implications of plate tectonics for the Cenozoic evolution of western North America. Geol. Soc. Am. Bull. 81, 3513-3536. Castillo, P.R., 2012. Adakite petrogenesis. Lithos 134-135, 304-316. Rock specimens from the Kitakami granitoids were partly offered by Chung, S.L., Chu, M.F., Ji, J., O'Reilly, S.Y., Pearson, N.J., Liu, D., Lee, T.Y., Lo, C.H., Satoshi Kanisawa, all the specimens from the Abukuma granitoids were 2009. The nature and timing of crustal thickening in Southern Tibet: Geochemical and zircon Hf isotopic constraints from postcollisional adakites. Tectonophysics 477, offered by Shunso Ishihara and Kazuaki Okamoto, and specimens from 36-48. the Nigata granitoides are offered also by Shunso Ishihara. Analysts Cole, R.B., Stewart, B.W., 2009. Continental margin volcanism at sites of spreading ridge other than authors are also shown in the appendix tables. subduction: Examples from southern Alaska and western California. Tectonophysics 464, 118-136. Cox, K.G., Bell, J.D., Pankhurst, R.J., 1979. The interpretation of igneous rocks. George, Allen and Unwin, London. Appendix A. Supplementary data Defant, M.J., Drummond, M.S., 1990. Derivation of some modern arc magmas by melting of young subducted lithosphere. Nature 347, 662-665. Defant, M.J., Jackson, T.E., Drummond, M.S., De Boer, J.Z., Bellon, H., Feigenson, M.D., Supplementary data to this article can be found online at https:// Maury, R.C., Stewart, R.H., 1992. The geochemistry of young volcanism throughout doi.org/10.1016/jseaes.2019.103968. western Panama and southeastern Costa Rica: an overview. J. Geol. Soc. Lond. 149, 13 S. Osozawa, et al. Jourmal of Asian Earth Sciences 184 (2019) 103968 569-579. complexes. Geology 22, 1135-1138. Deng, J., Xiaoyong Yang, X., Li, S., Gu, G., Mastoi, A.S., Sun, W., 2016. Partial melting of Osozawa, S., 1997a. The cessation of igneous activity and uplift when an actively subducted paleo-Pacific plate during the early Cretaceous: Constraint from adakitic spreading ridge is subducted beneath an island arc. The Island Arc. 6, 361-371. rocks in the Shaxi porphyry Cu-Au deposit, Lower Yangtze River Belt. Lithos 262, Osozawa, S., 1997b. Major transform duplexing along the eastern margin of Cretaceous 651-667. Eurasia. In: Flower, M. (Ed.), AGU Geodynamics Series, Mantle Dynamics and Plate Falloon, T.J., Danyushevsky, L.V., Crawford, A.J., Meffre, S., Woodhead, J.D., Bloomer, Interactions in East Asia. AGU, Washington, D.C, pp. 245-257. S.H., 2008. Boninites and adakites from the northern termination of the Tonga Osozawa, S., Pavlis, T., 2007. The high P/T Sambagawa extrusional wedge, Japan. J. Trench: Implications for adakite petrogenesis. J. Petrol. 49, 697-715. Struct. Geol. 29, 131-1147. Fiannacca, R., Cirrincione, R., Bonanno, F., Carciotto, M.M., 2015. Source-inherited Osozawa, S., Tsai, C.H., Wakabayashi, J., 2012. Folding of granite and Cretaceous ex- compositional diversity in granite batholiths: The geochemical message of Late humation associated with regional-scale flexural slip folding and ridge subduction, Paleozoic intrusive magmatism in central Calabria (southern taly). Lithos 236-237, Kitakami zone, northeast Japan. J. Asian Earth Sci. 59, 85-98. 123-140. Osozawa, S., Wakabayashi, J., 2012. Exhumation of Triassic HP-LT rocks by upright Frost, B.R., Barnes, C.G., Collins, W.J., Arculus, R.J., Ellis, D.J., Frost, C.D., 2001. A extrusional domes and overlying detachment faults, Ishigaki-jima, Ryukyu islands. J. geochemical classification for granitic rocks. J. Petrol. 42, 2033-2048. Asian Earth Sci. 59, 70-84. Fukunari, T., Wallis, S.R., 2007. Structural evidence for large-scale top-to-the north Osozawa, S., Wakabayashi, J., 2015. Late stage exhumation of HP metamorphic rocks, <0.01 progressive localization of strain, and changes in transport direction, Sambagawa Arc. 16, 243-261. belt, Japan. J. Struct. Geol. 75, 1-16. Furlong, K.P., 1984. Lithospheric behavior with triple junction migration: An example Osozawa, S., Wakabayashi, J., 2017. Variety of origins and exhumation histories of based on the Mendocino triple junction. Physics of Earth and Planetary Interiors 36, Sambagawa eclogite interpreted through the veil of extensive structural and meta- 213-223. morphic overprinting. In: Bianchini, G., Bodinier, J.L, Braga, R., Wilson, M. (Eds.). Hirahara, Y., et al., 2015. Space variation of Sr-Nd-Hf isotopic compositions in from The crust-mantle and lithosphere-asthenosphere boundaries: Insights from xenoliths, orogenic deep sections and geophysical studies. Geological Society of America Mineral. Petrol. Sci. 44, 91-111 In Japanese with Englich abstract. Special Paper, vol. 526, pp. 49-71. Ishihara, S., Orihashi, Y., 2015. Cretaceous granitoids and their zircon U-Pb ages across Rickwood, P.C., 1989. Boundary lines within petrologic diagrams which use oxides of the south-central part of the Abukuma Highland, Japan. Isl. Arc 24, 159-168. major and minor elements. Lithos 22, 247-263. Ishihara, S., Hamano, K., Ikegami, A., 1998. Isotopic evaluation on the genesis of the Streckeisen, A., Le Maitre, R.W.L., 1979. A chemical approximation to modal QAPF classfication of the ineous rocks: Neues Jahrbuch fur Mineralogie Abhandlungen Jahn, B.M., 2010. Accretionary orogen and evolution of the Japanese islands- implica- 136, 169-206. tions from a Sr-Nd isotopic study of the Phanerozoic granitoids from SW Japan. Am. Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: J. Sci. 310, 1210-1249. implications for mantle compositions and processes. In: Saunders, A.D., Norry, M.J. Jahn, B.M., Glikson, A.Y., Peucat, J.-J., Hickman, A.H., 1981. REE geochemistry and (Eds.), Magmatism in the Ocean Basins, Geological Society Special Publication isotopic data of Archaean silicic volcanics and granitoids from the Pilbara Block, London, vol. 42, pp. 313-345. western Australia: implications for the early crustal evolution. Geochim. Cosmochim. Takahashi, Y., Mikoshiba, M., Kubo, K., Iwano, H., Danhara, T., Hirata, T., 2016. Zircon Acta 45, 1633-1652. U-Pb ages of plutonic rocks in the southern Abukuma Mountains: Implications for Jahn, B.M., Usuki, M., Usuki, T., Chung, S.L., 2014. Generation of Cenozoic granitoids in Cretaceous geotectonic evolution of the Abukuma Belt. Isl. Arc 25, 154-188. Takasu, A,1989, P-Thistories of peridotite and amphibolite blocks in theSanbagawa geochemical analyses, and implications for crustal evolution. Am. J. Sci. 314, metamorphic belt, Japan. In: Daly, J.S., Cliff, R.A, Yardley, B.W.D. (Eds.), The 704-750. 0.12 Kanisawa, S., 1974. Granitic rocks closely associated with the lower Cretaceous volcanic vol. 43, Pp. 533-538. rocks in the Kitakami Mountains, northeast Japan. J. Geol. Soc. Japan 80, 355-367. Tokunaga, S., Nakai, S., Orihashi, Y., 2010. Two types of adakites revealed by Kano, K., Kurimoto, C., Iwaya, T, Hoshizumi, H., Matsuura, H., Makimoto, H., 2002. 238U-230Th disequilibrium from Daisen Volcano, southwestern Japan. Geochem. J. Geological Map of Japan, 5th ed. Geological Survey of Japan. 44, 379-386. Kay, R.W., 1978. Aleutian magnesian andesites: melts from subducted Pacific Ocean Tsuchiya, N., 2008. Petrogenesis of adakites and their geological significance. Earth Sci. crust. J. Volcanol. Geoth. Res. 4, 117-132. (Chikyu Kagaku) 62, 161-182. In Japanese with Englich abstract. Kay, S.M., Ramos, V.A., Marquez, M., 1993. Evidence in Cerro Pampa volcanic rocks of Tsuchiya, N., Kanisawa, S., 1994. Early Cretaceous Sr rich silicic magmatism by slab 58.4 melting in the Kitakami Mountains, northeast Japan. J. Geophys. Res. 9, 703-714. 22205-22220. Kon, Y., Ejima, T., Morita, S., Takagi, T., 2015. Spatial U-Pb age distribution of plutonic Tsuchiya, N., Kimura, J., Kagami, H., 2007. Petrogenesis of Early Cretaceous adakitic 0.02 granites from the Kitakami Mountains, Japan. J. Volcanol. Geoth. Res. 167, 134-159. 110, 145-149. Tsuchiya, N., Takeda, T., Adachi, T., Nakano, N., Osanai, Y., Adachi, Y., 2015. Early Kon, Y., Takagi, T., 2012. U-Pb zircon ages of Abukuma granitic rocks in the western Cretaceous adakitic magmatism and tectonics in the Kitakami Mountains, Japan. 7.94 Japan. Magaz. Mineral. Petrol. Sci. 44, 69-90. In Japanese with Englich abstract. Konig, S., Schuth, S., Munker, C., Qopoto, C., 2007. The role of slab melting in the petro- Utsunomiya, A., Jahn, B.M., Okamoto, K., Ota, T., Shinjoe, H., 2011. Intra-oceanic island genesis of high Mg andesites: evidence from Simbo volcano, Solomon Islands. arc origin for Iratsu eclogites of the Sanbagawa belt, central Shikoku, southwest Contrib. Miner. Petrol. 153, 85-103. Japan. Chem. Geol. 280, 97-114. Kubota, K, Takeshita, T., 2008. Paleocene large-scale normal faulting along the Median Wallis, S.R., 1998. Exhuming the Sanbagawa metamorphic belt: the importance of tec- Tectonic Line, western Shikoku, Japan. The Island Arcm 17, 129-151. tonic discontinuities. J. Metamorph. Geol. 16, 83-95. Martin, H., 1999. The adakitic magmas: modern analogues of Archaean granitoids. Lithos Whalen, J.B., Frost C., 2013. The Q-ANOR diagram: A tool for the petrogenetic and 46, 411-429. tectonomagmatic characterization of granitic suites. In: Conference: south-central Martin, H., Smithies, R.H., Rapp, R., Moyen, J.F., Champion, D., 2005. An overview of section, Geological Society of America, at Austin, Texas USA. vein Wallis, S., Anczkiewicz, R., Endo, S., Aoya, M., Platt, J.P., Thirlwall, M., Hirata, T., 2009. some implications for crustal evolution. Lithos 79, 1-24. Plate movements, ductile deformation and geochronology of the Sanbagawa belt, SW Mikoshiba, M.U., Kanisawa, S., 2008. Ptrochemical characteristics of the Tono plutonic Japan: Tectonic significance of 89-88 Ma Lu-Hf eclogite ages. J. Metamorph. Geol. complex, Kitakami Mountains. Earth Sci. (Chikyu Kagaku) 62, 183-201 (In Japanese 27, 93-105. with English abstract). Wilson, M., 1989. Igneous Petrogenesis, 466 pp. Unwin Hyman, London. Morris, P.A., 1995. Slab melting as an explanation of Quaternary volcanism and aseis- Yan, J., Liu, J., Li, Q., Xing, G., Liu, X., Xie, J., Chu, X, Chen, Z., 2015. In situ zircon Hf-O micity in Southwest Japan. Geology 23, 395-398. isotopic analyses of late Mesozoic magmatic rocks in the Lower Yangtze River Belt, Murata, M., Kanisawa, S., Ueda, Y., Takeda, N., 1974. Base of the Silurian System and the central eastern China: implications for petrogenesis and geodynamic evolution. pre-Silurian granites in the Kitakami Massif Northeast, Japan. J. Geol. Soc. Japan 80, Lithos 227, 57-76. 475-486 In Japanese with Englich abstract. Zellmer, G.F., Iizuka, Y., Miyoshi, M., Tamura, Y., Tatsumi, Y., 2012. Lower crustal H20 Okamoto, K., Shinjoe, H., Katayama, I., Terada, K., Sano, Y., 2004. SHRIMP U-Pb dating controls on the formation of adakitic melts. Geology 40, 487-490. of quartz-bearing eclogite from the Sanbagawa belt, south-west Japan: Implications Zhao, Z.H., Xiong, X.L., Wang, Q., Wyman, D.A., Bao, Z.W., Bai, Z.H., Qiao, Y.L, 2008. for metamorphic evolution of subducted protolith. Terra Nova 16, 81-89. Underplating-related adakites in Xinjiang Tianshan, China. Lithos 102, 374-391. Orihashi, Y., Ishihara, H., 2015. Finding of 115 Ma adakitic tonalite east of the Futaba Zhang, G., Niu, Y., Song, S., Zhang, L., Tian, Z., Christy, A.G., Han, L., 2015. Trace ele- Fault, Abukuma Highland, and its regional geological implications. J. Geol. Soc. ment behavior and P-T-t evolution during partial melting of exhumed eclogite in the Japan 121, 167-171. North Qaidam UHPM belt (NW China): Implications for adakite genesis. Lithos 226, Osozawa, S., 1994. Plate reconstruction based upon age data of Japanese accretionary 65-80.
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Journal of Volcanology and Geothermal Research, 29 (1986) 413--450 413
Elsevier Science Publishers B.V., Amsterdam -- Printed in The Netherlands
GEOCHEMISTRY OF THE QUATERNARY VOLCANIC ROCKS OF THE
NORTHEAST JAPAN ARC
M. SAKUYAMA* and R.W. NESBITT
Geology Department, University of Southampton, Southampton, Hants. S09 5NH, U.K.
(Received August 26, 1985;revised and accepted January 21, 1986)
ABSTRACT
Sakuyama, M. and Nesbitt, R.W., 1986. Geochemistry of the Quaternary volcanic rocks
of the Northeast Japan arc. In: I. Kushiro (Editor), M. Sakuyama and H. Fukuyama
Memorial Volume. J. Volcanol. Geotherm. Res., 29: 413--450.
Major and trace-element data are presented for a series of lavas from 17 volcanic
centres in the NE Japan arc. These represent a transect of the Quaternary arc from its
volcanic front (type A volcanoes) in the east, across its central zone (type B) to its
western margin (type C). Rocks range from basalt to dacite in composition and the
variation is attributed to fractionation of plagioclase, mafic silicates and titanomagnetite.
For comparative purposes trace~element data for each volcano are presented on a 55 wt.%
SiO2 normalised basis. The volcanoes display the characteristic features of arc volcanism
viz enrichment in large ion lithophile elements (Rb, Sr, Ba, K) and depletion in high
field strength elements (Zr, Nb, Hf, Ta). Using these data and derived primary magma
compositions (Tatsumi et al., 1983), mixing calculations give the percentages of fraction-
ating mineral phases necessary to derive the 55 wt.% SiO2 liquids. These data together
with published distribution coefficients allow the calculation of trace~element abundances
in the primary magmas. Reciprocal trace-element plots indicate that irrespective of
position with respect to the volcanic front, the elements, La, Ba, K, Sr, Nb, Y and Zr of
the primary magmas all lie on a common line strongly suggesting that the melts are the
products of varying degrees of melting of a common homogeneous source. However,
those melts close to the volcanic front (type A) show anomalous Rb, Th and Pb values
suggesting that their mantle source was enriched in these elements. Further calculations
suggest that the HFS depletion is not due to retention by residual mineral phases but is
a feature of the source. Thus the data suggest: (a) that the source of the arc volcanism
was homogeneous; and (b) that the characteristic chemical features were an inherent
feature of the source. Since Zr/Nb ratios in the source melts are close to those found
in N-type MORB it is suggested that a major component of arc-volcanic source rocks
was a depleted mantle of the type which has given melts of N-type MORB composition.
LIL elements added to such a depleted source are derived by the loss of silica-rich aque-
ous fluids from the descending (subducting) slab. Such losses occur within the first
100 km and are fixed in the overlying mantle wedge. This contaminated material is
transported down by drag-induced convection, producing a homogeneous, but meta-
somatised mantle within the melting zone. Diapiric uprise produced by small-scale melt-
ing, produces further melting, with the major control of melt chemistry being the degree
of partial melting. Superimposed on this chemistry is a localised addition of Pb, Rb
and Th in the source of those volcanoes nearest the volcanic front.
*deceased.
0377-0273/86/$03.50 © 1986 Elsevier Science Publishers B.V.
414
INTRODUCTION
Processes controlling the chemical composition of arc magmas fall into
two categories. The first group of processes involves the generation of
primary magmas in the upper mantle and the second involves the subsequent
history of the melts on their way to the surface. In the first category we
can include source rock composition, degree of partial melting and depth
of melt-residue separation whilst the second group includes crystal fraction-
ation, crustal contamination and magma mixing. This rather idealised model
does not however bring out one of the most important characteristics of
arc magmatism which was originally demonstrated by Kuno (1959) for the
Quaternary volcanics across Japan viz the relationship between subduction
processes and arc magmatism. Kuno's demonstration of a zonal arrangement
in the geochemistry of Japanese arc magmas was subsequently verified in
other arcs (e.g. Kuno, 1966; Dickinson and Hatherton, 1967; Nielson and
Stoiber, 1973) and although there are some obvious exceptions (see Arculus
and Johnson, 1978) these serve only to illustrate the complexity of the
process.
Although most arc volcanic assemblages are dominated by andesites,
it is clear that these are the result of fractional crystallisation of basaltic
liquids in which anhydrous phases (olivine, pyroxene, plagioclase and titano-
magnetite) are removed to produce a systematic enrichment in silica (e.g.
Gill, 1981). This process does not however explain certain chemical features
such as ratios among large ion lithophile (LIL) elements and in particular
the systematic differences observed between the geochemistry of mid-ocean
ridge basalts (MORB) and arc tholeiites (e.g. Perfit et al., 1980). It is this
difference which holds the key to the understanding of the role of sub-
duction processes in arc magmatism. For example, whilst it is now generally
believed a combination of processes such as mantle heterogeneity and/or
degree of partial melting is responsible for the LIL element variations of
arc rocks (e.g. Perfit et al., 1980; Saunders et al., 1980; Kay, 1980), there
is also abundant evidence that the down-going slab contributes in some
way to the variation. This latter view is strongly supported by isotopic
data (e.g. Armstrong, 1971; Church, 1976; DePaolo and Wasserburg, 1977;
Hawkesworth et al., 1979; Brown et al., 1982; Barreiro, 1983; Morris and
Hart, 1983), which points to an involvement of deep-sea sediments and
sea water in the mantle source of arc magmas. Indeed the general view
now seems to be that it is the introduction of water into the overlying
mantle wedge from the downgoing oceanic slab which provokes the vol-
canism.
We believe that the most plausible mechanisms in controlling the across-
arc geochemistry of some arcs involve variations in the composition of the
source and variations in the degree of partial melting. In order to assess
the importance of these mechanisms, it is necessary to separate the chemical
features attributable to secondary processes (crystal fractionations, crust--
415
mantle interaction etc) from those found in primary magmas. In this paper
we present major- and trace~lement data on Quaternary volcanic rocks
from the NE Japan arc and attempt to evaluate the data in terms of primary
processes such as source heterogeneity and degree of partial melting.
GEOLOGICAL SETTING AND GEOCHEMICAL OVERVIEW
The NE Japan volcanic arc is located on the convergent plate boundary
between the Eurasian and Pacific plates with a relative velocity of collision
of about 8.7--9.9 cm/yr {Minster et al., 1974) (Fig. 1). The leading edge
of the downgoing Pacific slab is considered to reach about 700 km depth
based on deep seismic activity (Utsu, 1974). Zonal arrangements of some
geophysical parameters across the arc, such as seismic structure of the
crust, gravity anomalies, heat flow and seismic activity are well summarised
by Yoshii {1979).
Igneous activity of the NE Japan arc has been continuous since the
early Miocene (Sugimura et al., 1963). The intense andesitic activity of the
early Miocene, about 23--15 Ma, was followed by the weak bimodal basalt-
dacite activity during the late Miocene to Pliocene. During the Quaternary,
volcanic activity (mainly andesitic) became intense again and all the vol-
-- 130~E 500KM 140°E
Volcanic front i/ Trench axis
~PAN
\ o, IPaclflc Plate 1 t : 35%1 !
Aklta
PACIFIC OCEAN 140°E .~¢~ ~,0 O" 144=E
s'"
4 / /
s,, . -- 7J/f'- ...
°,
I P..ippi..s.P,e~ ~ ~ 0 30oH (, !
• ..I.." • •
• ;:.....'. :'~
PHILIPPINE ~[\ " "" ¢
\. TOKY -- T /
200 KM
• TYPE A
(]) TYPE B
(~ TYPE C
Fig. 1. Position of NE Japan arc. Inset shows location of volcanoes and the symbols
the type of volcano. Dashed lines indicate depth to Wadati-Benioff zone.
416
canoes which can be recognised at present are the products of this latest
activity. The present active zone coincides with the Miocene volcanic region,
but with about 40 km westward retreat of the volcanic front. In the NE
Japan arc, there are about 100 Quaternary volcanic centres distributed
in a zone with a width of about 100 km extending from central Honshu
to the west of Hokkaido island. The distribution of volcanic materials is
-- asymmetric, exponentially decreasing toward the Sea of Japan from a
maximum at the volcanic front (Aramaki and Ui, 1982). The zonal arrange-
ment of the chemical composition of Quatermary volcanic rocks was first
pointed out by Kuno (1959). He recognised the increase in the alkali-lime
index toward the back-arc side and related it to deep seismic activity dipping
northwestward. This variation is largely the result of increasing alkali con-
tents, especially K20.
Recent trace~element studies show that incompatible elements such as
Rb, Ba, U, light REE and Sr, behave in a similar manner to K20 (e.g. Masuda
et al., 1975; Fukuyama, 1978; Katsui et al., 1978; Masuda, 1979). Lateral
variations of REE patterns such as increasing La/Sm, La/Yb and LIL/HFS
ratios, away from the volcanic front have also been demonstrated (Masuda,
1966, 1979; Masuda et al., 1975; Fujimaki and Kurasawa, 1980; Tatsumi
and Nakano, 1984). It has been suggested by Sakuyama (1979) and Ishikawa
et al. (1980), that H20 and F behave in a similar manner to K and other
incompatible elements.
Zonal arrangements can be seen also in isotope ratios. Tatsumoto and
Knight (1969) observed a decrease of 2°6pb/2°4pb, 2°Tpb/2°4pb and 2°aPb/
2°4Pb ratios, away from the volcanic front in NE Japan and suggested the
involvement of downgoing oceanic sediments in magma generation. A
decreasing ratio of aTSr/86Sr away from the volcanic front was pointed out
on one traverse of the arc, by Hedge and Knight (1969) and this was later
extended to the whole of the NE Japan arc by Notsu (1983). Nohda and
Wasserburg (1981) have, however, demonstrated the complexities of the
problem in their study of Nd and Sr isotopes in the volcanoes of NE Honshu.
Their data (3 samples) lie on a line parallel to the well-documented mantle
array with the most westerly volcano (Kampu) showing the lowest Sr and
Nd radiogenic values. TheY invoke the presence of a local spreading centre
in the Japan Sea and suggest that the resultant oceanic crust has influenced
the Quaternary volcanism on the west side of Honshu. In contrast, their
interpretation of a section from Hakone (south) to Myoko (north) involves
increasing contamination of mantle magmas by continental crustal material.
Sample description
At least 4 samples were analysed from 17 volcanic centres in an attempt
to cover as wide a ~range of major chemical variation as possible. Rocks
range from basalt to dacite with SiO2 variation from 48.7 to 70.2 wt.%.
Most of the analysed samples are free from magma-mixing textures, particu-
TABLE 1
Average modal abundances
Ol Aug Opx PI Mt Qz Hb Gm
Type A 1.5 2.1 3.3 21.3 0.8 2.0 1.9 71.9
(34) (45) (50) (53) (38) (7) (2) (53)
0.1--7.0 0.1-- 6.3 0.1-- 9.1 4.3--43.2 0.1--2.3 0.1--6.6 1.8--2.0 56.6--92.6
Type B 1.6 3.8 3.0 23.0 1.3 2.4 1.1 67.6
(10) (21) (20) (21) (21) (4) (5) (21)
0.2--3.8 0.6-- 7.8 0.5-- 6.4 2.9--33.7 0.3--2.5 0.9--5.4 0.1--3.0 58.2--94.7
Type C 5.9 6.6 4.4 29.0 1.1 -- 1.4 57.2
(5) (10) (6) (10) (9) (4) (10)
2.4--9.6 1.4--15.7 0.5--11.3 24.5--38.0 0.2--2.4 0.2--2.8 4.5--69.7
Ol : olivine; Aug: augite; Opx: orthopyroxene; P1: plagioclase ; Mt : titanomagnetite; Qz: quartz; Hb : hornblende ; Gm: groundmass.
Averages were taken only if each phase was present. Numbers in parentheses are the numbers of averaged samples.
418
larly dusty-zoned or resorbed plagioclase phenocrysts or disequilibrium
phenocryst assemblages such as Mg-rich olivine with quartz and augite
overgrowths on hypersthene phenocrysts. Eighteen samples out of 87 show
slight evidence of magma mixing, but even in these cases, textural evidence
suggests that mixing has occurred between magmas whose compositions
were very similar. Rocks containing hydrous phenocryst phases such as
hornblende and biotite were excluded as far as possible, since distribution
coefficients of some elements are so high between these phases and mag-
matic melt that their treatment by fractional crystallisation models becomes
complicated. The exceptions are 11 samples mainly from the back-arc region
(Oshima-Oshima, Iwaki, Kampu, Moriyoshi, Cholai and Gassan volcanoes)
which contain hornblende phenocrysts with a maximum modal abundance
of 3 vol.% in one rock from Moriyoshi volcano, but with an average mode
of 1.3 vol.%. 56 samples (64%) are augite-orthopyroxene andesite or basic
andesite, whose phenocrysts phases are plagioclase, augite, orthopyroxene
and titanomagnetite with or without olivine. Some of the andesites from
the volcanic front do not contain titanomagnetite phenocrysts. Phenocrysts
in olivine-augite basalt, the most common basalt type, are olivine, augite
and plagioclase with or without titanomagnetite. The modal abundance of
augite is greater in basalts from the back-arc region than in those from the
frontal arc. In some frontal volcanoes e.g. Iwate and Zao, orthopyroxene-
olivine basalts are observed. Some dacites contain quartz phenocrysts.
Ranges and the averages of the modal amounts of each phenocryst are
summarised from representative rock types in Table 1. The total phenocryst
contents range from 5.3 vol.% to 55 vol.% with an average of 31.1 vol.%
showing that most of the analysed samples are typical porphyritic arc
volcanic rocks (Ewart, 1976).
GEOCHEMISTRY
Analytical methods
Major elements were measured using a Telsec betaprobe, which is a
simultaneous X-ray fluorescence spectrometer employing a rastering electron
beam as the exciting source rather than an X-ray tube. The samples were
analysed as carbon-coated glass discs made up of 1 part sample to 5 parts
lithium tetraborate. Matrix corrections used the method described by
Lucas-Tooth and Price (1961). Trace elements were measured on pressed
powders by conventional X-ray fluorescence spectrometry (Phflips PW
1400) using standard correction techniques (e.g. Nesbitt and Stanley, 1980;
Tertian and Claisse, 1982). Eighty-seven samples were measured for Rb,
St, Zr, Y, V, Ni and Cr and from these, 34 samples were measured for Th,
Ba, La, Nb and Pb. Twenty replicate analyses of one sample gave the follow-
ing precision (one sigma standard deviation) Nb, 2.3 + 0.3; Y, 18 + 0.3;
Zr, 18 + 0.4; Sr, 109 + 0.8; Ni, 163 + 1; Rb, 2.0 + 0.5;Zn, 66.0+- .08;
419
Pb, 6.0 + 1; Th, 6.0 + 1 (all in ppm). Detection limits are 1 ppm for Rb,
Sr, Zr, Y, and Nb; 3 ppm for Th, Ni, Pb; 5 ppm for Cr, V and La and 12
ppm for Ba. Peak counting times of 100 s were used except for Nb where
300 s was used. Representative major- and trace-element chemical compo-
sitions are listed in Table 2 in which total iron is tabulated as FeO. Complete
analysed data is available from R.W.N.
Major-element chemistry
For the purposes of this study, the volcanoes are classified according
to their geographical position. Type A volcanoes are those on the volcanic
front, type B are those immediately behind and parallel to the front and
type C are those at least 50 km west of the front. The classification of
each volcano is shown on Table 2 and Fig. 1. Among the major elements,
A1203, FeO, K20 and TiO2 show the greatest variation between the various
types of volcano. In Fig. 2 the data from this study are plotted together
with previously published data of Kawano and Aoki (1960), Kawano et al.,
(1961); Katsui et al., (1978). A1203 in types B and C volcanoes has a maxi-
mum at about 56 wt.% SiO2 but there is no maximum in type A volcanoes.
On the other hand, FeO in type A rocks shows a maximum at about 55 wt.%
SiO2 but there is no maximum in type B and C rocks. The rate of decrease
of FeO against increasing SiO2 is highest in type C rocks. TiO2 variation
2o Type A
AI203 o• •
• "~" "'¥1: ""
." ..., ": • .'"
Wt' 15
10 i i * io i ~ q i
:~:~'" • F.O*
"1 $ ~ •
o• • • Type B
AI2103
I l" • • •
• ~oo •
", •2•. •
FeO*
~.... • ,. ".~,... •
• • ..e•• •
• " "e4" .
KzO K20
(C) (Cl
.... • ,
50 60 50 60
SlO2Wt % SiO2Wt % Type C
. AI203
..~, o'o" • • °e fi" :
15
n * ~a n
FeO* 10
• o •
o
• .°
. q i i i ~
K20
• , (C)
~lee" 2 ° • °*"
, , , L i i ~ i
50 60
SiO2Wt %
Fig. 2. Major-element (total Fe as FeO) concentrations, plotted against SiO~ wt.% for
the three types of volcano. Large circles with crosses represent normalised values at
55 wt.% (see text). Parallel lines in K:O diagram divide the field up into low, medium
and high-K andesites (after Gill, 1981).
b.~
TABLE 2
Representative whole-rock chemical compositions
Type A
Volcano: Osore Hakkoda Towada Hachirnantai Iwate Akita-koma Kurikoma Funagata Zao
Sample: OSO-27 HKD-11 TOW-3 HASP-l-1 IWT-12 HASP-10-6 KUR-13 FUN-7 ZAO-10
SiO~ (wt.%) 54.29 52.45 54.51 57.52 51.75 52.41 57.81 52.21 51.85
0,76 0.86 0.86 0.93 0.71 0.85 0.75 0.62 0.92
18.23 17.68 17.35 15.97 18.10 19.80 16.60 18.78 18.70
9.96 9.42 9.41 9.55 9.78 8.81 8.24 8.79 9.33
0.20 0.19 0.18 0.18 0.20 0.16 0.16 0.17 0.19
4.56 6.10 4.82 3.94 7.05 4.00 4.70 5.54 5.82
9.27 10.78 9.91 8,34 10.10 11.27 8.36 11.94 10.21
2.54 2.43 2.69 2.99 2.25 2.52 2.59 1.83 2.52
0.25 0.18 0.34 0.67 0.15 0.24 0.85 0.21 0.50
0.13 0.09 0.12 0.10 0.10 0.10 0.11 0.08 0.15 Ti O
AI,O3
FeO*
MnO
MgO
CaO
Na20
K20
P,Os
Pb(ppm) 7~ 6.0 8.7 6.9 5.8 6.0 8,5 4,6 4.0
Th 1.0 1.0 2.1 1.5 1.1 2.1 3.2 0.7 1.8
Rb 3 2 7 10 3 6 23 4.0 11
Ba 115 92 126 199 57 102 232 58 185
Sr 259 279 312 271 281 293 222 169 374
La 7.1 3.3 5.0 8,1 2,3 3.8 9.8 6.6 6.5
Nb 0.7 0.9 1.6 1.8 1.2 1.5 2.0 0.5 1.8
Zr 40 40 55 76 31 55 79 29 57
Y 26 22 24 28 15 24 25 17 21
V 246 281 276 292 235 257 207 245 279
Ni 13 26 17 7.2 43 9 17 16 32
Cr 136 83 152 7.9 97 41 187 76 92
TABLE 2 (continued)
Type B Type C
Volcano : Kayo Iwaki Moriyoshi Yakeyama Chokai Gassan Oshima-Oshima Kampu
Sample: HASP-12-5 IWK-4 HASP-7-4 HASP-5-3 CHO-36 GAS-26 OSOS-53 KMP-5
SiO2 (wt.%) 51.89 57.61 58.41 56.00 54.61 59.79 50.54 54.16
TiO2 0.83 0.83 0.98 0.82 0.89 0.83 0.91 0.75
Al203 21.12 17.38 16.81 18.09 16.56 16.83 17.21 18.49
FeO* 9.09 8.08 7.99 8.08 8,09 7.32 7.92 7,66
MnO 0.20 0.19 0.17 0.18 0.17 0.16 0.17 0.17
MgO 2.70 3,95 3.88 4.36 5.91 3.44 7.86 4.64
CaO 11.31 7.86 7.18 8.44 9.19 6.44 10.01 9.19
Na20 2.60 3.34 3.31 3.20 3.06 3.23 3.05 3.33
K20 0.35 0.82 1.27 0.79 1.52 1.93 2.12 1,53
P2 05 0.11 0.13 0.18 0.22 0.16 0.17 0.38 0.25
Pb(ppm) 5.9 10.8 11.1 9.0 8.0 12.6 7.0
Th 0,7 2.0 2,6 2.2 4.7 6,3 4.3
Rb 5 17 26 18 40 54 64
Ba 120 310 478 422 440 568 508
Sr 342 332 372 458 433 419 563
La 7.4 9.7 1.9 14.5 14.2 13.3 23.7
Nb 1.3 1.8 3.5 2.5 2.1 2.6 2.9
Zr 52 94 137 93 125 101 101
Y 26 32 38 24 30 24 24
V 273 178 200 166 229 180 248
Ni 16 15 16 52 7 100
Cr 42 141 40 155 4 189 10.3
8.2
47
868
680
26.2
2.4
100
22
209
30
197
FeO*: Total Fe recalculated as FeO.
bD
422
is similar to that of FeO. TiO2 contents of type B and C rocks monoto-
nously decrease against increasing SiO2, but that of type A rocks has its
maximum at about 58 wt.% SiO2. There is a small Na20 variation between
the three types; the level of Na20 contents seems to increase from type A
to type C. There is no conspicuous difference in CaO and MgO variation.
These differences in major-element variation can be accounted for by
different degrees of fractionation of plagioclase, mafic silicates and titano-
magnetite. A1203 variation suggests that the fractionation of plagioclase
is very effective in type A magma, but becomes effective only at the late
stage of fractional crystaUisation in the type B and C magmas. As shown by
Sakuyama (1979) the water content of the magmas across the NE Japan arc
increases towards the back-arc side. Increasing water content depresses
the stability field of plagioclase, resulting in less effective plagioclase frac-
tionation in the type C magma. On the other hand, FeO and TiO2 variations
can be ascribed to the effects of fractionation of titanomagnetite. This
mineral is very effectively subtracted in the early stages of fractionation
in the type C magma. However, in the type A magma, titanomagnetite
fractionation is not dominant until the magma differentiates to andesitic
composition. Type B magmas show an intermediate degree of titanomag-
netite fractionation. Variation of modal abundances of titanomagnetite
phenocrysts is consistent with this interpretation. K20 contents vary widely
3 T
1
K20
wt %
40
2O
Y pprn i i I
o o
i I I
I I I
200 Xx~
Zr ppm
I
510 SiO 2 wt % 60 100~
50f
Rbl
PPml
8001
600}
400~
200~-
Sr /
ppm 50 i i i
o
Si02 wt % 60
o Oshima-Oshima
x Chokai
• Hakkoda
Fig. 3. Trace-element and K20 concentrations plotted against SiO2 wt.% for three rep-
resentative volcanoes. Filled circles = type A; crosses = type B; and open circles = type
C. Large circle represents normalised trace-element value at 55 wt.% SiO2.
423
and range from low-K tholeiitic series of type A volcanic through calc-
alkaline to slightly alkaline series of type C (Fig. 2).
Trace-element chemistry
To illustrate the behaviour of some of the trace elements, the data from
three representative volcanoes were selected. Figure 3 is a plot of SiO2
against Y, Zr, Rb and Sr for Oshima-Oshima (type C), Chokai (type B)
and Hakkoda (type A) and illustrates the typical variation displayed by
the incompatible elements. For comparative purposes the figure also con-
tains K20. The most important feature illustrated by Fig. 3 is that each
volcano is characterised by a distinct chemical variation for each element.
Apart from Sr the abundances of incompatible elements monotonously
increase with increasing SiO 2. These variations strongly support the argument
that all the volcanic rocks from a single volcano are either co-magmatic
and fractional crystallisation has played a major role in generating the
chemical variation within them, or magma mixing has been a controlling
process. Since the samples were expressly chosen to avoid the problem
of mixing, we favour the first model.
TABLE 3
Normalised K20 and trace-element abundances at SiO 2 = 55 wt.%
K 2 O Pb Th Rb Ba Sr La Nb Zr Y
Type A
Osore 0.26 7.8 1.0 8 119 244 7.9 0.8 43 21
Hakkoda 0.51 9.4 1.6 16 169 267 6.1 1.4 68 23
Towada 0.45 8.8 1.7 11 144 289 6.2 1.8 62 23
Hachimantai 0.55 5.5 1.6 15 162 244 7.8 1.8 66 21
Iwate 0.36 11.2 1.5 12 123 261 8.3 2.1 56 20
Akita-koma 0.43 7.8 2.4 12 150 274 6.7 2.1 72 25
Kurikoma 0.61 6.4 2.2 17 179 229 6.2 1.4 57 16
Funagata 0.41 7.9 2.0 12 128 210 12.1 1.4 58 35
Zao 0.67 5.1 1.9 18 184 269 7.1 1.9 62 22
Type B
Kayo 0.54 7.5 1.1 14 174 318 11.6 2.1 79 31
Iwaki 0.67 11.3 2.0 15 281 320 9.8 1.5 85 27
Moriyoshi 1.07 11.3 2.6 21 386 395 13.8 3.0 118 32
Yakeyama 0.78 7.7 2.4 23 302 371 11.6 2.3 88 22
Chokai 1.65 9.2 6.2 46 490 456 18.2 2.6 143 32
Gassan 1.40 13.0 8.1 40 727 400 21.0 3.5 130 30
Type C
Oshima-Oshima 2.44 7.6 5.8 87 579 644 26.3 3.6 114 26
Kampu 1.63 12.5 10.5 53 1041 665 31.9 3.1 124 23
All values are in ppm except for K20 (wt.%).
424
For purposes of comparison, the hypothetical trace~lement abundances
of each volcano at 55% SiO2 have been derived by least-squares fitting of
the data on the SiO2 variation diagram (Table 3). To minimise extrapolation
errors, a median SiO2 value of 55% was selected for the normalisation
procedure. In the case of Th, Ba, La, Nb and Pb where only two samples
from each volcano were measured, the hypothetical abundances at 55%
SiO2 were derived by obtaining the average of their ratios to Zr and then
multiplying by the Zr value obtained for that volcano by least-squares
fitting. The rationale behind this treatment is that trace-element ratios
among those elements under consideration are constant within the observed
range of chemical variation. In fact most ratios are very similar in samples
from a single volcano. The reason for the normalisation by SiO2 will be
discussed later. Those parameters for Oshima-Oshima, Chokai and Hakkoda
volcanoes are plotted on Fig. 3 as large circles.
TRACE-ELEMENT ABUNDANCES IN PRIMARY MAGMAS
Arc volcanic rocks, in general, have undergone extensive fractional crystal-
lisation, as clearly exemplified by low Ni and Cr abundances. In this paper,
we estimate the trace~element abundances in the primary magma using
major-element chemistry and distribution coefficients taken from the
literature.
Tatsumi et al. (1983) have estimated three primary basaltic compositions
of the NE Japan arc. They used published primitive basalt compositions
whose FeO/MgO and SiO2 wt.% are less than 2 and 53, respectively, and
recalculated back to the primary compositions which could have equil-
ibrated with Fo89 olivine. The resultant compositions were averaged for
three groups and these correspond to the three types in this paper. They
examined the validity of these compositions from various viewpoints such as
the effect of plagioclase and pyroxene subtraction or accumulation, and
primary olivine composition (Fo89), and concluded that the estimated
primary compositions could not be strongly biased by these effects. In this
paper, we accept the three deduced compositions as typical primary major-
element compositions of NE Japan arc volcanics (Table 4).
We have attempted to estimate the trace-element characteristics of these
three primary basaltic magmas by a series of back-calculations which are
based on our own major- and trace~element data. Initially, we estimated
major-element compositions at 55 wt.% SiO2 for each volcano (by least-
squares fitting) and from those we estimated typical compositions for
each of the three major volcanic zones (Table 4). If these typical andesites
were derived from the primary basalts then this was probably achieved by
fractional crystallisation of phases such as olivine, augite, orthopyroxene,
plagioclase and titanomagnetite (primary phenocrysts observed in the
samples). In order to estimate the amount of each phase subtracted from
the primary composition least-squares mixing calculations were performed
TABLE 4
Assumed chemical compositions for mixing calculation 425
1 2 3 4 5 6 7 8 9 10 11
SiO 2 49.11 49.39 49.71 55.0 55.0 55.0 40.0 51.9 55.7 48.3 0.0
TiO 2 1.01 0.85 0.74 0.7 0.9 0.9 0.0 0.6 0.2 0.0 20.4
AI203 15.45 15.70 14.97 18.8 17.3 17.3 0.0 2.3 0.7 33.2 0.5
FeO* 9.42 9.76 10.57 7.2 8.5 9.7 14.1 9.5 13.6 0.0 78.6
MgO 11.59 12.05 13.03 4.2 4.8 4.5 45.8 14.5 28.1 0.0 0.5
CaO 9.66 9.43 9.00 8.7 9.0 9.5 0.0 18.9 1.4 16.1 0.0
Na20 2.54 2.33 1.56 3.3 3.1 2.8 0.0 0.2 0.1 2.4 0.0
K~O 1.09 0.34 0.28 2.4 1.7 0.5 0.0 0.0 0.0 0.1 0.0
1. Deduced major composition of primary alkali olivine basalt (AOB) (Tatsumi et ah,
1983).
2. Deduced major composition of primary high-alumina basalt (HAB) (Tatsumi et al.,
1983).
3. Deduced major composition of primary olivine tholeiite (OTB) (Tatsumi et al.,
1983).
4. Approximate major composition of Type C magma at SiO~ 55 wt.%.
5. Approximate major composition of Type B magma at SiO 2 55 wt.%.
6. Approximate major composition of Type A magma at SiO 2 55 wt.%.
7. Olivine (FOB0).
8. Augite (En43., , Wo40.7 , Fs16.0 ).
9. Orthopyroxene (EnT0).
10. Plagioclase (Ans0).
11. Titanomagnetite (UsP~0).
(Bryan et al., 1969) and it was assumed that the chemical compositions
of subtracted phases were constant. The assumed compositions are given
in Table 4. The results of this calculation are shown in Table 5. In the
calculation of type C, orthopyroxene was not included since small negative
numbers of orthopyroxene appeared if it was included. Variation of as-
sumed chemical compositions within a reasonable range causes only slight
differences to the results. For example, a 5% increase in Fo content produces
only a 1% decrease of subtracted amount of olivine. This variation has no
essential effect on the arguments. Sums of squares of residuals are 0.09--0.38
(Table 5), which are well below an acceptable upper limit, of 1.5 (Luhr
and Carmichael, 1980; Mann, 1983).
Using the data of Table 4 it is possible to calculate trace-element con-
centrations in the primary melts by the combined use of the results of the
mixing calculations (for each volcano-type) and recommended D values
from the literature. The following equation was used for calculation; Clx =
(1 - F) (D x - 1) C ° in which C ° is the initial concentration of element x,
Clx is the present concentration of element x, F is the solidified fraction
and Dx is bulk distribution coefficient. D values used for each element
and phase are listed in Table 6, which were adopted from the recommended
426
TABLE 5
Results of crystal fractionation modelling
F O1 Aug Opx Pl Mt R 2
Type A 35.4 11.5 6.6 17.6 25.7 3.1 0.09
Type B 42.2 13.7 7.4 8.2 25.5 3.3 0.38
Type C 49.5 16.6 13.1 -- 17.6 2.9 0.23
Estimated proportions of mineral phases subtracted from primary basalts to derive 55
wt.% SiO~ andesites. F = weight fraction of residual (55% SiO2) liquid and R: = sum of
squ~es of residuals between observed and calculated. Compositions of primary basalt
liquids and the mineral phases used in the modelling are given in Table 4.
TABLE6
Assumed D values for the calculation of trace-element abundances in the primary magma
O1 Aug Opx P1 Mt
K 0.01 0.02 0.01 0.11 0.01
Pb 0.01 0.30 0.35 0.03 1.00
Th 0.01 0.01 0.05 0.01 0.10
Rb 0.01 0.02 0.02 0.07 0.01
Ba 0.01 0.02 0.02 0.16 0.01
Sr 0.01 0.08 0.03 1.8 0.01
La 0.01 0.10 0.05 0.20 0.10
Nb 0.01 0.30 0.35 0.03 1.00
Zr 0.01 0.10 0.03 0.01 0.10
Y 0.01 0.50 0.20 0.03 0.20
Based on values given by Gill (1981).
values of Gill (1981). D values for Pb were assumed to be the same as those
for Nb since a preliminary estimate of the bulk D during fractional crystal-
lisation gives us a similar value to that of Nb.
Calculated bulk distribution coefficients axe less than 0.1 for K, Rb,
Ba, Zr and Th, 0.11--0.12 for La, 0.15--0.18 for Nb, 0.22--0.27 for Y and
0.66-0.81 for St. All these values are consistently lower than those obtained
from natural chemical variation (e.g., Mann, 1983) because bulk distribution
coefficients obtained here are valid for the derivation of andesite from the
primary basaltic magmas and hence contain a greater olivine effect than
in the natural basalt-andesite sequence.
Two points were assumed in the above procedure: (1) constant mineral-
melt distribution coefficients (D) and (2) constant proportions of sub-
tracted minerals. Strictly speaking, there is no doubt that the Nernst-type
distribution coefficients are dependent on T, P and the chemical compo-
sitions of melt and solid phases. However, actual chemical variations pro-
427
duced by fractional crystallisation in natural magmas are well explained
by constant bulk D's over a wide range of chemical variation unless the
fractionating phases change (Okamoto, 1979; Mann, 1983). This validates
the first assumption at least within the precision required here. The second
assumption is clearly inconsistent with the actual fractionation process
since minerals start to crystaUise sequentially, especially in the initial stage
of the process (Sakuyama, 1983a). If the D's are low enough for all the
relevant phases, the assumption does not affect the results. Rayleigh frac-
tional crystallisation models (Neumann et al., 1954; All~gre et al., 1977)
suggest that 0.2 difference of D between two elements causes only 15%
difference in their ratio even after 50% solidification. This is the same order
of magnitude of the precision of the normalizing procedure adopted pre-
viously. Therefore, for all elements except Sr, the validity of the second
assumption does not cause any serious difference to the results.
In order to assess the case of Sr, we compare the following two extreme
models in which D of Sr is assumed to be 0.1 to solid X and 2.0 to solid Y.
Model 1: Initial 20% solidification is performed by phase X and next 10%
by solid Y. Model 2: The same amounts of solid X and solid Y crystallise
TABLE 7
Estimated K and trace-element abundances in the primary magmas
K Pb Th Rb Ba Sr La Nb Zr Y
Type A
Osore 810 2.9 0.34 2.8 46 185 3.1 0.34 17 9.6
Hakkoda 1590 3.4 0.59 5.9 65 203 2.4 0.58 26 10.5
Towada 1400 3.2 0.63 4.2 55 219 2.4 0.76 24 10.5
Hachimantai 1710 2.0 0.56 5.6 62 185 3.1 0.75 25 9.7
Iwate 1120 4.1 0.56 4.4 47 198 3.3 0.90 22 9.2
Akita-koma. 1340 2.8 0.86 4.3 58 208 2.6 0.91 28 11.5
Kurikoma 1900 2.3 0.80 6.4 69 174 2.5 0.60 22 7.3
Funagata 1280 2.9 0.73 4.4 49 159 4.8 0.60 22 13.7
Zao 2080 1.9 0.69 6.6 71 204 2.8 0.81 24 10.1
Type B
Kayo 1690 3.2 0.45 5.0 78 241 5.4 1.03 30 14.2
Iwaki 2470 4.9 0.87 6.5 127 271 4.6 0.74 38 13.7
Moriyoshi 3940 4.9 1.13 9.0 174 335 6.4 1.46 53 16.3
Yakeyarna 2870 3.3 1.04 9.9 137 314 5.4 1.10 40 11.4
Chokai 6070 4.0 2.68 20.0 222 386 8.5 1.27 64 16.3
Gassan 5150 5.6 3.49 17.4 329 339 9.8 1.69 59 15.3
Type C
Oshima-Oshima 10370 3.8 2.92 44.1 300 506 14.0 1.98 60 15.6
Kampu 6930 6.2 5.26 26.6 539 522 17.0 1.70 66 13.8
All values in ppm.
428
until 30% crystallisation is complete. Enrichment factors of Sr (concen-
tration in the present liquid/that in the initial liquid) are calculated to be
1.07 for the Model 1 and 1.10 for Model 2. Thus, the second assumption
does not have any significant effect on the estimated initial Sr value and is
ignored in further discussion.
Estimated abundances of trace elements in the primary magmas for
respective volcanic centres are listed in Table 7, and some of them are
plotted in Fig. 4 normalised by the N-type MORB source composition
proposed by Wood et al. (1979a). These normalised patterns clearly show
the chemical characteristics typical of arc-type magmas such as a strong
enrichment of LIL elements and strong depltetion of HFS elements (e.g.,
Hawkesworth et al., 1979; Perfit et al., 1980). Another important point
to notice in Fig. 4 is the sub-parallel nature of each pattern, which sug-
gests a similar source for all NE Japan arc magmas. The first problem to
clarify is the principal mechanism for generating this sub-parallel variation
in the abundance patterns of the primary magmas.
~500 Ob
03
¥
~100
50
o
lO
±
> 5,
N
d <~
n-
O 1 Z I
i. ABUNDANCES IN PRIMARY MAGMAS
\ II- -OSHIMA'OSHIMA (Type C) I
L"... O--CHOKAI (Type B)
- -~--~,. I,. ~ IWAKI "
. "~ \ O--ZAO (Type A)
\ "B O- -TOWADA ,,
-~-"- - --eX X' E}- -OSORE ,,
, -..
', ~, /i I I_~," ',',\~ ',_'4 \~ 'I~" ,e\\
F[ 1 I I L J J I I--
Rb Ba Th K Nb La Sr Zr y
Fig. 4. Trace-element concentrations in primary magmas normalised to their concen-
trations in N-type MORB source. Data plotted are for representative volcanoes from the
three zones.
429
BATCH PARTIAL MELTING AND SOURCE COMPOSITION
Reciprocal plots
The variation in concentration of an element x during general non-modal
batch partial melting processes can be formulated using the following equa-
tion (Shaw, 1970):
C1/C° = 1/[D x + F(1-Px) ] (1) X / X
in which C ° = abundance of the element x in the initial mantle before
melting; Clx = abundance of the element x in the partial melt; D x = bulk
distribution coefficient for the initial starting assemblage; Px = average of
distribution coefficients of each mineral weighted by the mass fraction of
each mineral (Pi) in the melt and F = weight fraction of the melt. The
same equation can be drawn for an element y:
C1/C ° y. y = 1/[Dy + F(1-Py)] (2)
F can be eliminated from equations (1) and (2) and the following equation
is obtained (Minster and All~gre, 1978).
C°(1-Py)I 1 [D~ Dx(1-Py)
Cy (1-P x )
This equation implies that various melts produced by various degrees of
1 1 1/C 1 diagrams, batch partial melting give straight line variations on /C x -
(called 'reciprocal plots' hereafter), provided Dx, Dy, Px and Py are constant
during melting.
Reciprocal plots can be a useful tool for discriminating batch partial
melting from source heterogeneity caused by mixing of two components.
The latter mechanism produces a hyperbolic trend on reciprocal plots and
straight line variation on normal co-variance plots. Fractional melting will
make an exponential curve on normal plots (MacCaskie, 1984). Theoreti-
cally, this could be discriminated from batch partial melting on a reciprocal
plot, but the scatter of data points usually makes it difficult. In the case of
arc magmas, fractional melting can be neglected since much volcanological
evidence suggests the presence of magma storage systems beneath volcanoes.
This will result in the homogenisation of each successive melt resulting in
compositions similar to those produced by batch partial melting.
Another important advantage of the reciprocal plot is that, as shown by
eqn. (3), the slope and intercept of the straight line are determined by
initial abundances of elements and their distribution coefficients during
batch partial melting. Minster and All~gre (1978) used eqn. (3) to delineate
a batch partial melting model for interpreting the REE and Th abundances
in basanitoids and alkali basalts from Grenada, Lesser Antilles. Since garnet
430
plays an important role in controlling REE abundances in this province,
they could not make any approximation on Dx, Dy, Px and Py in eqn. (3).
But if the elements which will be considered are restricted to incompatible
elements whose D's are less than 0.5, it is possible to make the simplification
that D x and Dy are equal to Px and Py without any essential effect on the
results, giving the following equation:
1 [C~(1-Dy)] 1 Dy-D x
- -- + (4)
Cyl (l_Dx) Cxl Cy°(1-Dx)
The slope, Ay x and intercept, By x of eqn. (4) are expressed as:
C° (1-Dy)
Ay x = (5)
Cy(1-Dx)
Dy-D x
By x - (6)
c ,(1-Dx)
Although eqns. (5) and (6) cannot give us unique solutions, two variables
in these equations can be expressed as functions of another two variables.
Reciprocal plots of estimated trace-element abundances in the primary
magmas of the representative samples are plotted in Fig. 5. Two important
points appear in this figure: (1) type B and C compositions lie on a straight
t , , , , , 30 ~ 0
'° f . h L
I ~Ba O0
'°f i i 2O
lO ;,,, L i i
6O
40
0 0
2 • • 0
6 ' i , i0~ i , 20 r r i I I I i
10 • • 0
2
/ I 2JO I 410 I 60 I / 20 40 60 I
loyz, lo;Vz, 2 • 0 0
"°I , i i I i i i
10 2 0
10 • 0
8 • 0
lOyZr
Fig. 5. Reciprocal plots of several trace elements for NE Japan arc volcanics. Values
represent estimated compositions in the primary magmas. Open circles = type A; closed
circles = type B; squares = type C.
431
line for all pairs of element variations; (2) for type A composition, Ba, La,
K, St, Nb, Y and Zr plot on the same straight lines defined by the type B
and C compositions, but Th, Rb and Pb show significant deviations.
Straight line relationships on reciprocal plots of most element pairs
strongly suggests that varying degrees of batch partial melting (at constant
D) of a homogeneous source is a principal factor controlling the chemical
variation. As shown previously, simple two-component mixing in the source
will not show these relationships (at least with constant degree of partial
melting). An explanation involving source heterogeneity coupled with
varying degrees of partial melting might be possible, but it requires ad hoc
coupling between them.
In contrast to the straight line relationships of Ba, La, K, Sr, Nb, Y and
Zr through all the types, systematic deviations of Th, Rb and Pb of type A
compositions requires some other explanation. A lack of systematic changes
of crustal thickness of structure across the volcanic zone (Yoshii, 1979)
makes a model involving high level crustal contamination unlikely although
the isotope data of Nohda and Wasserburg (1981) argues against such a
view. There are two other possibilities: (1) D of Th, Rb and Pb for the
generation of type A compositions are different from those of type B and
C with similar source compositions; (2) D was constant during melting,
but concentrations of Th, Rb and Pb in the source mantle were higher in
the type A source than in the type B and C sources.
In order to examine the first possibility, we compare the bulk D of Th
in the type A and B compositions. Starting from the same composition,
a sudden change of the bulk D of Th at some point during increasing de-
grees of melting will result in a break in straight line variation to another
straight line with different slope and intercept. The two lines must be
continuous, and this requires a smaller slope and higher intercept for the
type A compositions on the reciprocal 1/Th - 1/Zr plot in Fig. 5, than for
type B compositions. Using eqns. (5) and (6), this can be expressed as
follows:
C~r (I-DTh)] < [C~r(I-DTh)]
Sl°pe: [ ~TTh~JA ~JB
intercept: [ DTh-DZr 1 >[ DTh- Dzr ]
C~r)"A "C~h (1-Dzr)~ S
Since Czr , CTh and DZr are constant on both sides, the equation reduces
to the following relationship:
[DTh]A > [DTh]B
Thus, in order to interpret the data according to the first model, D of Th
must be higher in the type A composition than in the type B composition.
432
The same argument can be drawn for Rb whilst the data are less convincing
for Pb. As shown previously, other chemical data show that the degree of
melting is higher in type A than in type B compositions, so that Th, Rb
(and perhaps Pb) must be retained by some phases with increasing degrees of
partial melting. This is very unlikely to take place in the normal course
of peridotitic melting. We conclude that the second possibility (initial
source enrichment) is therefore the most plausible. Average enrichment
factors in the source which produced type A magmas are estimated about
50% for Rb and Th, and about 25% for Pb.
Source compositions and bulk D
As pointed out previously, estimated trace-element abundances in primary
magmas show typical arc characteristics, such as high LIL/HFS elemental
ratios (Fig. 4). With respect to the origin of these characteristics, two inter-
pretations have been proposed. (1) Strong depletion of HFS, such as Zr,
Nb, Ta and Hf is caused by the retention of these elements in a Ti-bearing
phase (sphene, rutile, etc.) during partial melting processes (Hellman and
Green, 1979; Green, 1981; Morris and Hart, 1983). In this model, the
source composition is assumed to be similar to MORB-type or intra-oceanic
island basalt type sources. (2) Trace-element patterns essentially reflect the
chemical features of the source mantle and Ti-bearing phases play only
a minor role. This problem can be approached directly by estimating initial
source compositions from reciprocal plots. As a first step, the slope (Ayx)
and intercept (Byx) of the straight line on reciprocal plots were determined
for all possible pairs of elements by least-squares approximation. For those
pairs which involve Th, Rb and Pb, least-squares approximation was per-
formed on type B and C compositions only. 36 sets of Ay x and By x were
obtained for 9 elements, and were solved analytically using CSrO as a param-
eter and assuming DTh = 0.001. From examination of Byx for all pairs
of elements, the following order of bulk D is readily obtained [see eqn.
(6)]:
DTh <DRb ~DK ~DBa <DLa<DZr~DNb ~Dpb <DSr<Dy
Th is the most incompatible element of the nine. The diopside-melt distri-
bution coefficient for Th is less than 0.007 in the Di-An-Ab system (see
review by Irving, 1978). This means that bulk D during partial melting
in peridotitic upper mantle would be less than 0.001. However, based on
experimental data, Hanson (1977) estimated bulk D of Rb during partial
melting at less than 0.001. Therefore, the assumed value of D for Th (0.001),
may be the maximum value possible. Fortunately, this assumption does
not have any effect on estimation of C~x and C ° from eqns. (5) and (6) • . . Y
since these can be reduced to the foUowmg equation free from D:
o = By x X o X o C °- AyxX Cy Cy C x (7)
433
Furthermore, changing DTh has no major effect on estimated D of the
other elements. For example, if DTh is reduced to 0.0001 there is only a
9% difference for Rb and less than 5% difference for other elements.
Having obtained values for Ay x and Byx, the following calculation proce-
dures were followed to derive initial source compositions:
(1) C ° was calculated for each element from eqn. (7) using assumed
C~r and linear parameters A and B, which were determined on element
pairs including Sr.
(2) C ° values for each element were used for the estimation of C ° values
of other elements, again through eqn. (7), using A and B parameters for the
element pair. Therefore, each element has eight different estimated C °
values.
(3) These different values were simply averaged for each element in order
to eliminate some possible bias from the particular pair of elements. The
range and average for each of the elements under consideration are given in
Fig. 6 for the case where C~r = 50 ppm.
Th
Rb , At Sr = 50 ppm
, ['1 O.05_O,01ppm
0.!
r~ ~ ~-l 0.49_+0.08 ppm
i
0.5
120 160 200 240 188±34ppm
i
Ba
La
Zr
Pb I I ' ~ ~ ~ 1 8 g~2+O ppm
0 10
0.5 044±O.09ppm
~-~ , 43~0.7ppm
5
A .... /7 ~] ~,'~ I~ 0.60±0.15ppm
0 05
Nb
J J
0.1 0.2
n 3, O000m
0 5
Fig. 6. Histograms of estimated source compositions for several trace elements and K
(see text). The figure shows the range and average values where C~r = 50 ppm.
434
(4) D of each element was estimated two ways. The first, from parameters
A and B for that element and Th, using estimated initial concentrations
of each element and assumed DTh. The other is from two sets of the param-
eters including three elements: viz. Th, the element under consideration
and one other element which has lower D than the element under consider-
ation. For example, Zr has 5 different estimated values which come from Th,
Rb, K, Ba and La. Resultant values were averaged to eliminate possible bias.
Figure 7 shows the range and average value based on DTh = 0.001.
It is important to note that all of these calculations are based on the
assumption that we are dealing with melt-solid equilibria only. Our justi-
fication is that the results are internally consistent within the observed data
and in our view give support to the batch partial-melting model.
Bulk D
0-01 0.1
9 ISr
~ IP b
4~Zr
ql La
<1 K
--~1 Rb Dfh=O-O01
0"1
~4 Ba lallrlllll
DSr
0-01
~ so
ec 40
"o~
~ so
u~
~ o 20
tu 10
%0
5
E
o. 4 o.
>" 3
T. 2
¢
- 1
0 I i i t I I I I I i
~ ~ ~ ~ ~ ~ ~ Type
~ ~ Type B j~
I I I I I I I I I I
50 100
Initial Sr ppm
Fig. 7. Estimated bulk D for several trace elements and K (see text), where DTh = 0.001
and C~r = 50 ppm. The figure shows the range of estimated values for each element.
Average values indicated by solid triangle.
Fig. 8. a. Variation of D strontium against its initial source concentration, b. Degree
of partial melting against initial source concentration of Sr. c. Y abundance in the source
against Sr in the source. The figure suggests that Sr concentration in the source can be
constrained to 40--50 ppm.
435
In the calculations, the choice of a value for C~r is clearly important.
We have used the following constraints to fix the C~r at between 40 and 60
ppm:
(1) During partial melting of a peridotitic upper mantle, DSr greater than
0.1 is unlikely, based on previous experimental results (e.g., Irving, 1978).
Hence from the variation of DSr against C~r shown in Fig. 8a, C~r would
be expected to be about 40 ppm.
(2) Expected degrees of partial melting can be calculated for primary
magmas based on initial trace element abundances as functions of C~r.
Calculated results are shown in Fig. 8b. As previously discussed, major-
element features of the type A primary compositions are characteristic
of olivine tholeiite. Experimental work (Green, 1973; Jaques and Green,
1980; Takahashi et al., 1981) suggests that olivine tholeiitic melts can be
produced by 20--30% partial melting of peridotitic upper mantle. This
gives C~r between 50 and 70 ppm (Fig. 8b).
(3) Wood et al. (1979a) estimated Y abundances in 'primordial mantle'
and 'N-type MORB source' as 4.9 and 4.1 ppm, respectively. During metaso-
matic enrichment processes in the upper mantle, Y does not seem to move
extensively (Wood et al., 1979b). As will be shown later, the source mantle
for NE Japan arc is likely to be a depleted N-type source mantle which has
been later enriched by LIL elements. Y is not extensively enriched in this
process and therefore we estimate Y abundance in the source to be less
than 4 ppm, so that C~r would be less than 60 ppm (Fig. 8c).
None of these constraints is tight enough to determine C~r value defi-
nitely, However, taking all these constraints into consideration, 40 to 50
TABLE 8
Internally consistent source abundances, bulk D, and possible variation of degree of
partial melting
o . 40 ppm 50 ppm 60 ppm CSr.
C ° (ppm) D x Cx ° (ppm) D x C ° (ppm) D x
Th 0.04 0.001 a 0.05 0.001 a 0.06 0.001 a
Rb 0.38 0.011 0.49 0.014 0.61 0.017
K 147 0.016 188 0.020 231 0.024
Ba 7.0 0.016 8.9 0.020 10.9 0.025
La 0.34 0.029 0.44 0.037 0.54 0.044
Zr 3.4 0.046 4.3 0.058 5.3 0.071
Pb 0.48 0.062 0.60 0.079 0.72 0.096
Nb 0.14 0.064 0.18 0.080 0.21 0.097
Sr 40 a 0.102 50 a 0.128 60 a 0.153
Y 2.6 0.18 3.2 0.22 3.7 0.26
Degree
of melting 1.1%--17.9% 1.3% -- 22.9% 1.6% -- 28.3%
aAssumed value.
436
ppm can be chosen as the preferred range of C~r. Internally consistent
source compositions, bulk distribution coefficients and expected degree
of partial melting at C~r = 40, 50 and 60 ppm are listed in Table 8, and
initial abundances at C~r = 5- ppm which are normalised to N-type MORB
source (Wood et al., 1979a) are plotted in Fig. 9.
This figure shows a striking parallelism between the estimated source
composition and the patterns of the observed volcanics. This suggests that
the source of the volcanics had already the distinctive depletion (of HFS)-
enrichment (of LIL) patterns prior to partial melting and that partial melting
processes were not involved in dictating the characteristic trace-element
500
100
0
== lo
0
w e~
t-
Z
m
a
_1
.,j
E
o z I0
0.1
0"0!
I I = I I I I l J 0.001
Rb Ba Th K Nb La Sr Zr Y ~ Mg
t
F.'tt Mo
?
/ \
I
\ Bulk D
Sr=5Oppm
DTh=O'O01
Ca
~Na
\\
11+)
04 06 08 tO 12 14 16
Ionic Radius A
Fig. 9. Range of estimated trace~lement concentrations (normalised to N-type MORB
source) in the primary magmas of the NE Japan arc. Also shown is estimated pattern
for their source mantle.
Fig. 10. Bulk D plotted against ionic radius (in A) for major and trace elements in arc
volcanic rocks. Estimates in bulk D are based on values of DTh = 0.001 and C ° for Sr =
50 ppm (see text for further explanation).
437
patterns of these arc volcanics. Hence models involving retention of HFS
elements by phases present in the source during partial melting are not
favoured. The implications of this conclusion will be discussed later.
Bulk D
Bulk D's are plotted in Fig. 10 against their ionic radius (Shannon and
Prewitt, 1969, 1970). Although the present data are not enough to draw
a complete curve, some reasonable assumptions allow us to improve it.
According to previous work (Masuda et al., 1975; Fujimaki and Kurasawa,
1980), basic rocks in NE Japan arc show no fractionation within the heavy
REE, and this is also confirmed by our results (Sakuyama and Nesbitt, un-
publ. data). Based on this, the curve for trivalent cations near Y is drawn
nearly flat. D of Sc which is trivalent and has an ionic radius of 0.75 A
between A1 and Y, was assumed to be greater than 0.5. According to ex-
perimental work, D of Sc is 0.3--0.4 for olivine, 0.5--1.4 for subcalcic
pyroxene and 1--100 for calcic pyroxene (see review by Irving, 1978).
The above assumption is a conservative one.
In addition to trace elements, roughly estimated D values for major
elements (Ti, A1, Fe, Mn, Mg, Ca and Na) are also plotted in Fig. 10. These
estimates were based on the following four assumptions:
(1) The major-element composition of the source is the average composi-
tion on 384 spinel lherzolite bulk chemical data compiled by Maaloe and
Aoki (1977) which include 301 continental and 83 oceanic spinel lherzo-
lites. The assumed composition is SiO~=44.2, A1203=2.05, FEO*=8.29,
MGO=42.21, CAO=1.92, Na20=0.27, TIO2=0.13, MnO=0.13. MnO abundance
in the assumed primary magmas which was not included in Table 4 was
taken from the original estimate by Tatsumi et al. (1983).
(2) Typical degrees of partial melting for type A, B and C primary com-
positions (Table 4) are assumed to be 20%, 10% and 5%, respectively,
based on the previous arguments.
(3) During partial melting, bulk D are constant even for major elements.
(4) D values are calculated from eqn. (2) assuming modal melting (Px =
Dx).
Diagrams of this type were initially proposed by Onuma et al. (1968)
to advocate crystal lattice controls on the mineral-melt distribution of
elements. The idea was later confirmed by Jensen (1973) and Matsui et al.
(1977). According to them, distribution coefficients of all elements with
the same ionic charge must plot on a smooth curve on the diagram which
may or may not show maximum or minimum points, and each curve is
sub-parallel. Ionic radii for these maxima correspond to the preferred ionic
radii of a mineral for a particular group of elements with the same ionic
charge. As clearly shown in Fig. 10, estimated bulk D's for elements with
the same valency produce sub-parallel smooth curves. This strongly supports
the validity of our estimates of D which are based on the batch partial
melting model.
438
The pattern obtained in Fig. 10, for example a sirnple peak at 0.7 A for
divalent and trivalent cations, is very similar to the pattern shown by bronzite
(Onuma et al., 1968; Matsui et al., 1977). On the other hand, characteristic
features of the D pattern for calcic clinopyroxene (the single peak at 0.9 A
for example) are absent. This suggests that within the source of the NE
Japan arc magmas, orthopyroxene was the dominant pyroxene -- a con-
clusion which is consistent with the generally accepted view of mantle
mineralogy (e.g., Carter, 1970; Maaloe and Aoki, 1977).
DISCUSSION
Origin of across-arc variation
It is our view that the data presented here strongly support varying
degrees of batch partial melting of a homogeneous source mantle as the
principal process producing chemical variation within arc primary magmas.
We are aware that such an interpretation does not match with published
isotope data on Sr and Nd (Nohda and Wasserburg, 1981). However, because
there appears to be sufficient ambiguity in the interpretation of these iso-
topes (particularly with respect to the role and nature of the continental
crust), we should avoid forcing our data into models which cannot be
sustained. Consequently, we believe that the increasing abundances of
incompatible elements in volcanic rocks toward the back-arc side imply a
decrease in degree of partial melting in the same direction. This is consistent
with a previous interpretation of H20 concentrations in the magmas (Sa-
kuyama, 1979) and Sr-Ba systematics (Onuma et al., 1983).
There are two possible ways of varying degrees of partial melting: (1)
varying temperature at constant pressure; and (2) varying pressure at con-
stant temperature. As summarised by Sakuyama (1983a), arc magmas have
a dual imprint in terms of the depth of generation. In one case shallow
depths are inferred both from the fairly low equilibration pressure, (< 25
kbar), of primary basaltic magma with mantle peridotite (e.g., Green and
Ringwood, 1967; Takahashi et al., 1981; Tatsumi et al., 1983) and a crustal
thickness control on the degree of partial melting (Sakuyama, 1983a, b).
On the other hand, there is a clear association of arc magmas with the
Wadati-Benioff zone, in that there is a rather constant depth of the zone
beneath the volcanic front [124 + 38 km, according to Gill (1981)]. Sim-
ilarly, there is a close correlation between the width of the volcanic zone
and the dip of the Benioff zone (Marsh, 1979; Shimozuru, 1981). Both
of these features imply a fundamental high-pressure control on the mag-
mas. Constraints on H20 concentration in the primary magmas, probably
less than 2.5 wt.% (Sakuyama, 1983b), requires unusually high temperatures,
1200--1300°C, at the top of the upper mantle. In order to reconcile the
conflicting evidence, Sakuyama (1983b) suggested that diapiric uprise of
the deep hot mantle from the vicinity of the Wadati-Benioff zone, and its
439
associated decompression, is the principal mechanism for arc-magma gener-
ation.
Furthermore, Tatsumi et al. (1983) showed that there is no across-arc
systematic difference in the temperature of the diapir itself. This means
that the depth at which the diapir ceases to rise is the factor controlling
the degree of partial melting. From consideration of the magma segregation
mechanism, they concluded that the mantle diapirs have been forced to
stop rising by an external factor. They suggested that a drastic change in
the viscosity of the surrounding upper mantle, which is controlled by its
thermal structure, would be a principal factor.
If we accept the diapiric uprise model, our data require the trace-element
composition for each diapir to be homogeneous except for Th, Rb and Pb.
This means that there is no extensive chemical exchange between the diapirs
and the surrounding upper mantle during their ascent, and also that the
chemical features of the original portion of the upper mantle from which
the diapirs started to rise is also homogeneous across the arc (except for
Th, Rb and Pb).
Geochemical features of source mantle for arc magmas
Chemical features of the source mantle for NE Japan arc magmas can be
summarised as follows:
(1) Depletion in terms of absolute abundances of HFS elements such as
Nb and Zr compared with N-type MORB source (Wood et al., 1979a, Sun
et al., 1979).
(2) Enrichment of LIL elements such as Rb, Ba, Th, K, La and Sr com-
pared with N-type MORB source.
(3) The depletion and enrichment of HFS and LIL elements (see 1 and 2)
are a homogeneous feature of the source region.
(4) The Zr/Nb ratio of the NE Japan source material (about 24, Table 8)
is very similar to N-type MORB (about 30) whereas other major basalt
provinces have ratios about 10 or less (Basaltic Volcanism Study Project,
1981). Change~ in this ratio during partial melting or fractionation are very
small because both Zr and Nb have a similar D (Pearce and Norry, 1979).
Therefore these ratios are representative of the source composition. We
therefore conclude that a major mantle component of the source of the arc
volcanics was a type which had previously provided N-type MORB i.e. it
had experienced a previous melting event. In Fig. 11, available Nb/Zr ratios
of arc volcanic rocks from the circum-Pacific region are plotted. (The data
source is available from RWN). The figure suggests that the N-type MORB
source mantle proposed for NE Japan can be also valid for other arcs.
Wood et al. (1979b) used a Th-Hf-Ta diagram which can discriminate N-type
MORB, E-type MORB, within-plate basalts and arc basalts. In their compiled
data, Ta/Hf ratios of arc basalts are very similar to those of N-type MORB,
about 7--30. Ta probably behaves like Nb (Wood et al., 1979b), so this
also supports our argument that an N-type MORB source is involved.
440
20
10 N--type MORB J f
Primordial Mantle
~/ OIB
T-type MORB
R Continental rift
Flood basalt
........ FT,~ ,~ -~
1 2 3 4 5 6 7 8 9 10 11
100 X Nb / Zr
Fig. 11. Nb/Zr ratios for basalts found in major tectonic provinces.
(5) Over-abundances of Rb, Th and Pb are characteristic only of the
type-A source and are superimposed on the homogeneous source.
If we accept the depleted N-type MORB source as the source mantle of
arc-magmas, LIL elements such as Rb, Ba, Th, K, La and Sr must have
been added to the depleted mantle (see Fig. 9). As pointed out by previous
workers (e.g., Perfit et al., 1980), the enrichment pattern is common to all
arcs and thus to all arc mantle sources. This means that the enrichment
process is likely to have characteristically taken place under arcs, strongly
suggesting chemical input from the subducted lithosphere.
Stern (1981) and Morris and Hart (1983) pointed out similar St-, Nd-
and Pb-isotopic ratios and Rb-K-Ba-Sr ratios between arc basalts (IAB)
and oceanic island basalt (OIB), and advocated a common mantle source
between them. The apparent depletion of HFS elements in IAB was at-
tributed to their retention by Ti-bearing phase during partial melting under
hydrous condition. However, in our view, the present study indicates that
the HFS depletion cannot be accounted for by retention during melting so
that it must be a chemical feature of the source mantle. This view is also
supported by the data of White and Patchett (1984) who showed that the
systematic overabundance of Cs, '°Be anomaly (Brown et al., 1982) and
excess :°TPb (Sun, 1980) in island arc basalt (IAB) is not consistent with the
deduced common source between IAB and OIB, but strongly suggests a
chemical input from the subduction zone.
Two distinct mechanisms have been considered for chemical input from
the subduction zone (see review by Ringwood, 1977; Wyllie and Sekine,
1982): (1) by an hydrous siliceous melt from the subducted oceanic crust;
and (2) by an aqueoUs fluid derived from dehydration of the subducted
441
oceanic crust and mantle. Both of these mechanisms have inherent problems.
The possible consequences of the hybridisation by hydrous siliceous
melts, originally proposed by Nicholls and Ringwood (1973), have been
extensively investigated theoretically and experimentally by Sekine and
Wyllie, (1982a,b,c). They demonstrated that phlogopite, quartz, garnet,
enstatite and jadeitic augite precipitate by hybridisation, resulting in phlog-
opite-bearing garnet pyroxenite which may produce a heterogeneous layer
above a subduction zone.
This process is not consistent with the geochemical data presented here,
which requires that the hydrous siliceous melt must be almost free of Zr
and Nb. This melt is supposed to be derived from quartz eclogite which has
been transformed from subducted oceanic crust under amphibolite facies
conditions. According to Pearce and Norry (1979), mineral-melt D of Nb
and Zr for clinopyroxene and garnet in a basic melt are less than 0.3, so
that the resultant melt will be enriched in these elements. It is possible to
assume the retention of these elements by some Ti-bearing phases under
hydrous condition. Rutile is the most possible candidate at greater than
30 kbar (Green, 1981) but its probable affinity for light REE (Green,
1981) will produce light REE depletion in the melt, which is opposite to
the observed enrichment pattern.
The second process, enrichment by aqueous fluid, is consistent with the
pattern of enrichment. HFS elements are insoluble in the aqueous fluid and
remain in solid phases, whilst the LIL elements appear to be selectively
taken up by the fluid (Tatsumi et al., 1986, this volume). This contrasting
element behaviour during dehydration of the subducting slab has also been
considered by others to be responsible for the chemical features of arc
volcanic rocks (e.g., Hawkesworth et al., 1979; Saunders et al., 1980).
The problem with this model is the disparity between where the volcanism
occurs in relation to the subducting slab and the depth where dehydratic.n
ceases. In the NE Japan arc, the depth of the Wadati-Benioff zone cor-
responding to the recent volcanic zone is 130--200 km (Utsu, 1974) and
according to Gill (1981), the average value is 124 + 38 km from other
arcs. However, Delaney and Helgeson (1978) suggest that the principal
hydrous phases of the subducting slab would have dehydrated before the
slab reached 100 km in depth. In other words, the depth of the main de-
hydration zone is too shallow to produce the initial partial melting of the
mantle which results in the volcanism.
One possible source of water below 100 km is phlogopite (Delaney and
Hegelson, 1978) but this would appear to offer only a small contribution,
because its abundance will be limited by the availability of K. We estimate
that phlogopite would contribute a maximum of 0.05 wt.% H20 even if
all the K in the descending slab (about 0.1 wt.% K20) was held in that
mineral. This is a maximum estimate because large amounts of K will be
carried out by an aqueous fluid during dehydration. Therefore, we cannot
assume a large contribution of H20 from the subducting slab below 100 km.
442
Another problem arises concerning the composition of the source of
arc magmas. Most models assume that the source is the mantle wedge over-
lying the descending oceanic slab and that magma generation occurs at
depths ranging from 130 to 200 km. Our previous discussion particularly
on the basis of HFS elements, suggests that one component in the source
was mantle of a type which gave rise to N-type MORB. It is difficult to
envisage such a component extending down to depths of 200 km. Indeed,
Ringwood (1982) considered the maximum depth of 'depleted pyrolite'
to be 80 km. Thus the problem is how to get such a source to depths below
100 km, metasomatise it (with an aqueous fluid) and then produce a homo-
geneous source region over a pressure range of 40--70 kbar. It is because of
these lines of argument that we cannot accept previously proposed mecha-
nisms in their simple form. The important parameter which is missing,
is a mechanism to produce homogenisation of N-type MORB source and the
added component (in the form of an aqueous fluid) prior to the final uprise
of the diapirs. Finally, it is important to stress the over-abundance of Rb,
Th and Pb in the A-type source. Again we favour an aqueous fluid input
from the subduction zone (because of the LIL element enrichment). How-
ever, the critical factor is that this input was superimposed on an already
homogenised mantle.
Magma generation model beneath arcs
In our view, the most feasible mechanism for arc-magma genesis involves
release of an aqueous fluid into the overlying mantle wedge from the de-
scending oceanic slab. We have already pointed out the discrepancy between
the initiation of volcanism and depth of dehydration and propose that this
can be overcome if the hybridised mantle wedge is dragged down by the
movement of the descending slab (Fig. 12). This mechanism would also
explain how 'depleted pyrolite' could be involved in magma genesis at
depth in excess of 100 km. ToksSz and Hsui (1978) showed that induced
convection in the wedge mantle is a necessary consequence of slab sub-
duction. The overall scale of the induced convection is difficult to estimate
but in our view downward flow of the mantle immediately adjacent to the
subduction zone is highly likely. Sekine and Wyllie (1982b, 1983) have
already suggested this as a mechanism to carry down the metasomatised
mantle, although in their model, metasomatism involved an hydrous sili-
ceous melt.
In our model, dehydration of subducted oceanic crust takes place con-
tinuously from several kbar to about 30 kbar (Delaney and Helgeson, 1978).
Expelled water migrates upward carrying LIL elements from the slab to
the overlying mantle. During the dehydration reaction, Zr, Nb, Hf, Ta and
probably Ti will be fixed in the downgoing slab within various phases or at
grain boundaries. The aqueous fluid probably contains large amounts of
silicate components, probably 5--20 wt.% (Boettcher and Wyllie, 1968;
500 400 300 200 100 0 KM 443
Fig. 12. Model for genesis of arc magmas based on loss of an aqueous fluid from the
descending slab and homogenisation of the metasomatised overlying mantle wedge by
slab-parallel forced convection.
Nakamura and Kushiro, 1974; Ryabchikov and Boettcher, 1980), in ad-
dition to the LIL elements. (Supporting evidence for this process comes
from the dominance of quartz-free eclogite xenoliths in kimberlites (Ma-
thias et al., 1970), which may result from the transformation of silica-
depleted oceanic crust transforming into quartz-free eclogite during sub-
duction.)
Our previous discussion indicates that a principal component of the
overlying mantle wedge is a depleted mantle produced as a result of the
extraction of N-type MORB at oceanic spreading ridges. This may corre-
spond to the 'depleted pyrolite' at the bottom of the oceanic plate proposed
by Ringwood (1982). Geochemical features of this mantle will be very
similar to an N-type MORB source but trace element abundances, including
both LIL and HFS elements, are low. Siliceous aqueous fluid which enters
into the peridotitic mantle will be fixed mainly by amphibole with sub-
ordinate phlogopite (Wyllie and Sekine, 1982). Excess silica may slightly
increase the volume of orthopyroxene by consuming olivine but this effect
will be negligible. (Sakuyama (1983b) estimated the possible maximum H20
abundance in the arc mantle to be at most 0.5 wt.%. If all the water has been
added during this metasomatism and this water contains even 50 wt.%
SiO2, the expected change of SiO2 concentration in the source mantle is
only 0.2 wt.%.)
444
During this metasomatism, the main geochemical features of the arc-
magma source material are developed, but the enrichment may be hetero-
geneous and some systematic dip-ward differences in composition might be
expected from the different dehydration reactions in the slab under in-
creasing temperature and pressure. However, induced convection in the
overlying wedge mantle will not only carry the hybridised zone down
into deeper levels, but also homogenise the heterogeneities in it by velocity
gradients in the flow. In addition, incipient partial melting will strongly
reduce the bulk viscosity of the mantle and induce shear zones in the hybrid-
ised mantle. This will further promote chemical homogenisation.
The upper pressure stability limit of amphibole in ultramafic compositions
is several kilobars higher than that in basaltic compositions, being stable
down to 100 km (Wyllie, 1979). Therefore, the downgoing metasomatised
mantle can transport water to deeper levels than can basaltic oceanic crust.
Even so, the breakdown depth is still less than the actual depth of the
Wadati-Benioff zone beneath the volcanic front (140 km in NE Japan arc).
This discrepancy can be explained if the initiation of melting is controlled
by the temperature distribution around the descending slab. Such a view
is supported both by the variation of depth of the Wadati-Benioff zone to
volcanic front in different arcs (90--200 km) and the crude correlation
between this distance and the age of the lithosphere (Fig. 13). In this model,
older, cooler subducting plates lower the temperature of the metasomatised
mantle wedge along the subduction zone resulting in deeper initiation of
the incipient melting. The positive correlation between rate of descent,
maximum depth of continuous subduction and age of descending plate
(Ruff and Kanamori, 1980) may enhance this cooling effect.
c 200
o
g
~ 150
c
o
N
0
~ 100
~K~
o I • I t
• me
50 100
Age of Subductlng Plates I
150 Ma
Fig. 13. Depth to the Benioff zone beneath volcanic front against the age of the sub-
ducting plates for several plate margins.
445
Once incipient partial melting commences the resultant decrease in densi-
ty caused by partial melting and concentrated shear movement results in
gravitational instability of the layer and eventual diapiric uprise of the less
viscous metasomatised mantle. Thus our model involves: (1) dehydration
of the downgoing slab in the first 100 km; (2) the resultant siliceous aqueous
fluid is fixed in amphibole in the overlying mantle wedge; (3) prior to meta-
somatism the overlying wedge has the overall composition of depleted
pyrolite (i.e. depleted in both LIL and HFS elements); (4) homogenisation
of the mantle wedge is achieved by downward convection parallel to the
subducting slab which carries the hybridised material into the melting zone;
(5) initiation of melting is controlled by the temperature profile around
the descending slab and this in turn may be a function of the age of the
slab; (6) melting of the overlying wedge initiates diapiric uprise due to
density instability and promotes further homogenisation by localising
shear zones parallel to the downgoing slab.
There still remains the problem of the selective enrichment of Rb, Pb and
Th in the type A source. We commented earlier that this enrichment was
superimposed on an already homogenised LIL-enriched source and that
in our view the enrichment was due to aqueous fluids. Although the con-
straints are clear, the explanation is not. We suspect that minor dehydration
reactions occurring beyond the critical 100 km limit may have produced
contamination in the already homogenised but metasomatised mantle
wedge. We propose that a rather localised event is responsible for this re-
striction to the type-A source.
ACKNOWLEDGMENTS
The authors thank R.J. Arculus, I.W. Croudace, J.B. Gill, J.C. Claoue-
Long and Y. Tatsumi for their helpful comments on the manuscript. Mrs.
A. Dunkley and Mr. B. Marsh assisted with the illustration and Mrs. L.
Emery typed several versions of the manuscript. Mr. K. Gosden assisted
in the analytical work. The work was supported by a grant to R.W.N. from
the Natural Environment Research Council (GR3/4437). R.W.N. acknowl-
edges his gratitude to his late colleague for two years of friendship and
stimulating research.
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Sakuyama (1986) geochemistry of volcanic rocks of northeast japan arc.txt
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Quaternary volcanic front at the junction of the South-west
Japan Arc and the Ryukyu Arc
Hiroki Kamata*
Osaka Center, Geological Survey of Japan, Gov. Bldg. No. 2, 4-1-67, Otemae, Chuo-ku,
Osaka 540, Japan
(Received 6November 1996; Accepted 5June 1997)
Abstract —The Hohi volcanic zone, a 70 by 40 km volcanic graben on Kyushu Island at the junc-
tion of the South-west Japan Arc and the Ryukyu Arc, has been formed since 6 Ma to the pre-sent. Quaternary volcanism, producing centers along a NE–SW trending line, the volcanic front,
started at about 1.5 Ma. This volcanic front crosses the E–W alignment that was built by the
Pliocene volcanoes of the Hohi volcanic zone. The Quaternary volcanic front is parallel to thedepth contours of the present deep-seismic plane, and is perpendicular to the current direction of
subduction of the Philippine Sea plate. The age of formation of the volcanic front coincides with
the time when several other tectonic events occurred, such as the counter-clockwise shift fromNNW to WNW in relative convergence direction of the Philippine Sea plate. The volcanic rockserupted after 1.5 Ma are more widely distributed, with a larger total volume in the denser seismic
zone than in the aseismic zone. This suggests that the volcanism on the volcanic front is closely
related to seismicity at about 120 km depth. The active seismic area may have accelerated to gen-erate a partially molten zone beyond the dragged hydrated peridotite layer above the downgoingslab in comparison to the aseismic area. #1998 Published by Elsevier Science Ltd. All rights
reserved
Introduction
The volcanic front is defined as the trenchward limit
of the alignment of volcanoes on a certain interval spa-
cing in typical subduction zones (Sugimura 1960), andgenerally forms a line parallel to the trench axis. Theformation of a volcanic front has been explained as aboundary of physical properties in the crust where
magmas can reach the surface (e.g. Sugimura and
Uyeda 1973). In most subduction zones, the volcanicfront typically lies about 110 km above the dippingseismic plane (Tatsumi 1986). Many island arcs show
that the volcanic front emerged abruptly at a certain
time, probably due to tectonic setting of each arc(Kamata 1987). The relationship in time and spaceduring the formation of the volcanic front, however,has not been well documented, mainly because of lack
of detailed, on-land geological, chronological and geo-
physical data for ancient volcanism.
Recent studies of Kyushu Island and Chogoku
District, located at the junction of the South-west
Japan Arc and the Ryukyu Arc (Fig. 1), have revealed
that the Quaternary volcanism in this area startedalong the NE–SW trending volcanic front. The age offormation of the volcanic front coincides with severalother tectonic events when the motion of Philippine
Sea plate changed its direction. In this paper I will
describe the chronological sequences and space re-lationships of the volcanism on Kyushu Island and inthe Chugoku District, and discuss the linkage between
generation of the Quaternary volcanic front and deepseismic activity.
Volcanism in Kyushu Island
In the central part of Kyushu Island, Plio-
Pleistocene volcanic rocks are widely distributed in a
70 by 40 km area. This area forms a volcano-tectonic
depression, the Hohi volcanic zone (HVZ in Fig. 2),and has been filled with more than 5000 km3of ande-
sitic material (Kamata 1989b). Geologic evidence,
including the E–W trend of normal faults that devel-oped in the Early to Late Quaternary, indicates thatthe HVZ was formed by N–S extension in Plio-
Pleistocene time (Figs 2 and 3) (Ikeda 1979, Chida and
Ikeda 1991). The HVZ forms a box-shaped graben sur-rounded by steep gravity gradients. Radiometric ages
of volcanic rocks in the HVZ show a systematic zona-
tion (Kamata 1989b) younging inwards from the mar-gins, with outer ages between 6 and 5 Ma and inner
ages of 0.2 Ma (Fig. 4). A similar zonal pattern is also
observed in the Bouguer gravity anomalies. No radio-metric ages older than 2 Ma were reported from within
theÿ5 mgal contour line, while those younger than 1
Ma were found only within the ÿ15 mgal contour line
(Fig. 4). Moreover, those rocks younger than 0.5 Maoccur in the subcircular region with an anomaly of
more thanÿ30 mgal (Ss in Fig. 4, Shishimuta Caldera)
(Kamata 1989a, Kamata et al. 1994). These features
imply close spatial and temporal links between the
Pliocene volcanism and the E–W elongated graben for-Journal of Asian Earth Sciences, Vol. 16, No. 1, pp. 67–75, 1998
#1998 Published by Elsevier Science Ltd. All rights reserved
Printed in Great Britain
1367-9120/98 $19.00 + 0.00
PII: S0743-9547(97)00044-5
* Present address: School of Earth Sciences, Faculty of
Integrated Human Studies, Kyoto University, Kyoto 606-8501, Japan.
67
Fig. 1. Index map of the Japanese Islands and associated plates and arcs.
Fig. 2. Dextral faults (the Oita–Kumamoto Tectonic Line and the Median Tectonic Line) and the Hohi volcanic zone
(HVZ; solid dots) in Kyushu and Shikoku Islands (Kamata 1992). Open arrows with PHS show the present subductiondirection of the Philippine Sea plate, after Seno (1977). Small solid arrows are active strike-slip faults. KT, Kokura–
Tagawa Fault Zone. Map location shown in Fig. 1.H. Kamata 68
mation. Pliocene volcanic rocks were mainly erupted
through E–W trending fissures (Kamata 1989b). The
eruption rate of the HVZ decreased with time. This is
indicated quantitatively by changes of both area anderupted volume for each 1 m.y. interval; early onduring the interval 5–4 Ma, a maximum activity isshown, producing about 3000 km
3(Kamata 1989b).
In contrast to the E–W trending Pliocene volcanism,
the volcanoes of the Quaternary volcanic frontemerged as a NE–SW trending chain within the HVZ.The volcanoes younger than 0.1 Ma, such as Yufu–Tsurumi, Kuju and Aso volcanoes, are aligned on this
chain (Fig. 4), which is oblique to E–W alignment
composed by the Pliocene volcanoes in the HVZ.Around these volcanoes occur three clusters of morethan 20 lava domes with a few small-scale stratovolca-noes (Kamata et al. 1988a), all of which are younger
than about 1 Ma (Fig. 5). These clusters are alsoformed in the vicinities of vent calderas for large-scale
ignimbrites which erupted more than 50 km
3magma,
such as the Aso, the Yabakei, the Imaichi and theYufugawa ignimbrites (Kamata 1989b, Hoshizumi andKamata 1991, Kamata et al. 1994). Thus, the chrono-
logical and geological data indicate that each cluster ofvolcanoes on a NE–SW alignment e/C128used more than100 km
3magma since about 1 Ma. According to
Nakada and Kamata (1991), the time and space re-lationships of these volcanic activities suggest that the
formation of the clusters can be ascribed to the ascent
of a mantle diapir beneath the volcanic front.
Within the HVZ, the volcanism on the volcanic
front can be traced back to 1 Ma. The time of com-
mencement of volcanicity at the volcanic front, how-ever, is dicult to determine in this area, because the
products of lower Pleistocene volcanism in the HVZ
have been buried and overprinted by upper Pleistocenevolcanic rocks. Therefore the commencement age offormation of the volcanic front has to be determined
outside the HVZ.
Volcanism in Chugoku District
Volcanism on the volcanic front in the HVZ is
linked to volcanoes outside the HVZ. They are ofnearly the same age, about 1.5 Ma, and part of the
same chain. North-east of the HVZ, they are rep-
resented by the small-scale lava-dome volcanoes ofFutagoyama and Himeshima [Fig. 6(A) and 6(B)].
Further northeast, the chain of volcanoes extends to
Fig. 3. Generalized geologic map of the Hohi volcanic zone (HVZ) after Kamata (1987) and Kamata (1989b). Map lo-
cation shown in Fig. 1.Quaternary Volcanic Front at the Junction of the South-west Japan and Ryukyu Arcs 69
the similarly small-scale lava-dome volcanoes of
Shikuma, Sengoku, Tokuyama–Kimpo and Aonoyamain the Chugoku District (Ono et al. 1981) [Fig. 6(C)].
These volcanoes are composed of hornblende ande-
site–dacite lava domes and cones and show the typical
features of independent groups of monogenetic volca-
noes (e.g. Nakamura 1977). In their lithological, mor-phological and geochemical characteristics, these
volcanoes closely resemble the andesite–dacite lava
domes and cones erupted within the HVZ (Kamata
1987, Kamata et al. 1988a). I propose that all of these
volcanoes on the NE–SW alignment emerged from the
same tectonic origin as the contemporaneous volcanoes
erupted on the volcanic front. The commencement ageof the volcanic activity at volcanic front has been
determined by K-Ar dating of the volcanoes located
outside the HVZ (Kaneoka and Suzuki 1970, Kamata
1987, Kamata et al. 1988b). Andesite lava domes and
cones at Futagoyama volcano were dated at 1.46–1.10Ma [Fig. 6(A)]. Dacite lava domes and flows at
Himeshima volcano were dated at 0.34–0.2 Ma
[Fig. 6(B)]. Andesite lava domes and cones of the
Chugoku District were dated at 1.28–0.23 Ma
[Fig. 6(C)]. Thus, the volcanoes on the chain outsidethe HVZ were emplaced after about 1.5 Ma. This
suggests that the NE–SW trending alignment of volca-
noes on the volcanic front originated at about 1.5 Ma
throughout the South-west Japan Arc.Tectonic implications
The NE–SW trending volcanoes on the volcanic
front are calc-alkaline and high-alkali tholeiitic seriesin composition (Nakada and Kamata 1991). They are
typical products of island-arc volcanism. The NE–SW
alignment of these volcanoes is parallel to the depthcontour of the deep-seismic plane (e.g. Yoshii andKobayashi 1981), and perpendicular to the currentsubducting direction (WNW) of the Philippine Sea(PHS) plate (Matsubara and Seno 1980, Seno andMaruyama 1984) (Fig. 7). Volcanoes on the volcanic
front in Kyushu Island show the characteristic across-
arc variation in mineral assemblages and whole-rockchemistry (Kamata et al. 1988a). Biotite–rhyolite lava
flows that erupted between 1 and 0.3 Ma occur to thenorthwest, the back-arc side, whereas hornblende-andesite and pyroxene-andesite lava domes occur onthe frontal-arc side (Fig. 5). This is consistent with the
lateral variation of phenocryst assemblages across typi-
cal island arcs, which has been accounted for by thevariation of H
2O content (e.g. Sakuyama 1979).
Therefore, the Quaternary volcanism on the volcanicfront is interpreted to have been generated by subduc-tion of the PHS plate. Hereafter I will focus on the re-lationship between the volcanism and the movement of
the PHS plate in the Quaternary time.
Fig. 4. Distribution of volcanic rocks and their radiometric ages superimposed on Bouguer anomaly map of the Hohi
volcanic zone (HVZ) after Kamata (1989b). Bouguer anomaly contours in mgal (after Komazawa and Kamata 1985,
Kamata 1993). Mi, Miyanoharu; Ss, Shishimuta. Map location shown in Fig. 1.H. Kamata 70
Matsuda (1980) and Kaizuka (1984) proposed that
the PHS plate was subducting northward at about 3Ma on the basis of structural changes in the north ofIzu Peninsula and the southern Kanto District inCentral Japan. They estimated that the PHS plate haschanged its subduction into a more westward direction
between 3 Ma and the present. The transitional period,
however, was not clarified due to lack of geologicalrecords during this period. Based on the analysis ofthe morphology in the triple junction o/C128 BosoPeninsula in Central Japan, as revealed by theSeabeam survey, Seno (1985) proposed that the PHS
plate changed its motion from NNW to WNW at
about 1.5 Ma.
The commencement of volcanic activity at the volca-
nic front (1.5 Ma) coincides with the age of the change
in direction of motion of the PHS plate proposed bySeno (1985). The following evidence corroborates thesignificant tectonic change at the junction of theSouth-west Japan and Ryukyu Arcs at about 1.5 Ma:(a) The active dextral motion on the Oita–KumamotoTectonic Line, which is the western extension of the
Median Tectonic Line, on the south margin of theHVZ (Fig. 2) (Kamata 1992) intensified at about 1.5Ma (Itoh et al. 1997). (b) Crustal rotation and rifting
have taken place during the past 2 m.y. in South
Kyushu and northern Okinawa trough (Kodama et al.
1995). (c) Calcalkalic high-Mg basaltic andesites wereerupted in the early stage of the HVZ activity (5–3Ma), whereas high-alumina basalts typical of island-arcs erupted in the later stage (2–0 Ma) (Nakada and
Kamata 1991). (d) The K
2O content of volcanic rocks
in Central Kyushu has increased after 1.6 Ma(Kamata 1989b). (e) Volcanic rocks around theKagoshima graben (Fig. 7) have been dated by K-Armethod at <1.5 Ma, suggesting that the volcanic front
on the northern Ryukyu Arc formed similarly at about
1.5 Ma (Sakaguchi 1988). Therefore, the generation ofthe volcanic front may have been closely related to thetectonic change caused by the westward transition ofthe convergence direction of the PHS plate.
Fig. 5. Regional distribution of volcanic rocks younger than 1 Ma around Shishimuta caldera, the vent for the Yabakei
and Imaichi pyroclastic flows (circle) (Kamata et al. 1994). Star, vent location of the Yufugawa pyroclastic flow.
Triangle, summit of mountain. Map location shown in Fig. 2.Quaternary Volcanic Front at the Junction of the South-west Japan and Ryukyu Arcs 71
Fig. 6. (A) Geologic map of Futagoyama volcano and K-Ar ages of volcanic rocks in Ma, after Moriyama et al. (1983),
Matsumoto and Narishige (1985), Kamata (1987) and Kamata et al. (1988b). Ft, Futagoyama; Ot, Otakeyama; In,
Inagawayama; Hb, Hebigatake; Ok, Okutaisan; Ak, Akane; Mo, Monjuyama. Map location shown in Fig. 2. (B)
Geologic map of Himeshima volcano and K-Ar ages of volcanic rocks in Ma, after Kasama and Huzita (1955),
Kaneoka and Suzuki (1970), Kamata (1987) and Kamata et al. (1988b). Map location shown in Fig. 2. (C) Geologic
map of the western part of Chugoku District and K-Ar ages of volcanic rocks in Ma, after Murakami (1975), Kamata(1987) and Kamata et al. (1988b). To, Tokuyama; Ta, Takeyama; Si, Shikumadake; se, Sengokudake; Ki, Kimposan; Cj,
Chojagahara; No, Nosakayama; Ku, Kumoimine; Ao, Aoyasannkita; Ck, Chikuratouge; Na, Nabeyamashima; Ab,
Abu. Map location shown in Fig. 7.
The volcanoes erupted after 1.5 Ma are located at
about 120 km above the deep seismic plane (Fig. 7).
The deep-seismic plane is detected under KyushuIsland south of Himeshima and under Shikoku island,
but undetected under the Chugoku District north of
Himeshima (Fig. 7). Shiono (1974) proposed that the
leading edge of the downgoing PHS slab has reached
the upper mantle in the seismic region (dotted line inFig. 7). On the basis of the ScSp observations, how-
ever, Nakanishi (1980) proposed the existence of an
aseismic continuation of the PHS slab beneath theChugoku District. He pointed out that the inclinedScS-to-P conversion interface may correspond to the
boundary between asthenosphere and an aseismic deadslab which had descended from the Nankai Trough.The seismic features are consistent with the pattern oferupted volume of Quaternary volcanism. The totalerupted volume after 1.5 Ma on the volcanic front out-side the HVZ (north of Futagoyama) is much less thanthat inside the HVZ (between Yufu–Tsurumi volcanoand Aso volcano) (Fig. 7). The boundaries between
the large-volume area and the small-volume area corre-
spond to those of seismic and aseismic regions (Fig. 7).
Fig. 7. Distribution of volcanic rocks, plate motions and the depth contour of deep-seismic plane at the junction of the
South-west Japan and Ryukyu Arcs. Black solid area shows the distribution of volcanic rocks younger than 1.5 Ma.Stippled area of angular shape represents the distribution of the Hohi volcanic zone (HVZ). Solid arrow shows the pre-
sent motion of the Philippine Sea plate after Seno (1977). Open arrow shows the past motion of the Philippine Sea plate
before 1.5 Ma according to Seno (1985). Dotted line is contour line of deep-seismic plane on the underthrusting plate(Yoshii and Kobayashi 1981). A, Aso volcano; K, Kuju volcano; Y, Yufu-Tsurumi volcano; a, Futagoyama volcano; b,
Himeshima volcano; c, Takeyama volcano; d, Shikuma volcano; e, Sengoku volcano; f, Tokuyama-Kimpo volcano; g,
Chojagahara volcano; h, Nosakayama volcano; i, Kumoimine volcano; j, Aoyasan volcano; k, Chikura volcano; l,Nabeyama volcano; p, Omine volcano; q, Akai volcano; r, Kumamoto-Kimbo volcano; S, Unzen volcano. Map location
shown in Fig. 1.Quaternary Volcanic Front at the Junction of the South-west Japan and Ryukyu Arcs 73
This suggests that the erupted volume in the past was
closely related to the present deep seismicity.
Experimental petrological data suggest that mantle
diapirs stop rising to segregate a primary magma in
subduction zones (e.g. Tatsumi et al. 1983). High-
pressure experimental data and thermodynamic calcu-lation indicate that dehydration of hydrous minerals in
the downdragged peridotite layer above the slab takes
place just beneath the volcanic front (e.g. Tatsumi andEggins 1995). Nakada and Kamata (1991) suggeststhat the mantle diapir under Kyushu Island was con-
taminated by a slab-derived component and segregated
a primary magma feeding e/C128usive lavas (Fig. 8).Greater amounts of volcanic materials are distributed
in the area above the denser seismic zone within the
dipping Wadati–Benio/C128 zone than those above lessdense seismic zone (Figs 7 and 8). The higher rate of
frictional/shear heating, which is suggested by denser
seismicity, may have caused a higher rate of, or greateramounts of magma production (e.g. Acharya 1981).
Although there was possibly a time gap between pre-
sent seismicity and Quarternary magmatism, it is likelythat the frictional/shear heating of the undergoing slabmay have been larger in the active seismic area in the
past as well as at present, resulting in accelerating the
partially molten zone above the dragged hydrated peri-dotite layer in comparison to the aseismic area.
Acknowledgements —Valuable comments and discussions for
this study by Kozo Uto, Tetsuzo Seno, Yoshiyuki Tatsumi,
Hideo Hoshizumi, Takehiro Koyaguchi, Keiichi Sakaguchi,
Kazuto Kodama, Setsuya Nakada, Keiji Takemura, YasutoItoh, Keiko Suzuki-Kamata and Kees Linthout are greatly
appreciated.
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kamata 1998 quaternary volcanic front ajapan Arc and Ryukyu Arc 1998.txt
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Journal of Asian Earth Sciences 72(2013) 75-87 Contents lists available at SciVerse ScienceDirect Journal of Asian Earth Sciences ELSEVIER journal homepage: www.elsevier.com/locate/jseaes Geology and tectonics of Japanese islands: A review - The ossMark key to understanding the geology of Asia Koji Wakita Geological Survey of Japan,AIST,1-1-1Higashi Tsukuba, Ibaraki 305-8567,Japan ＡRＴICＬＥ INＦＯ ABSTRACT Article history: a n n s e n s o no a a Available online 24 May 2012 ment is, however, expected, as the provenance of by detrital clasts of conglomerate, detrital zircons of metamorphic and sedimentary rocks, and as metamorphic rocks intruded by 500 Ma granites. Although Keywords: rocks of Paleozoic age are not widely distributed, rocks and formations of late Mesozoic to Cenozoic can Japan be found easily throughout Japan. Rocks of Jurassic age occur mainly in the Jurassic accretionary com- Japanese Islands plexes, which comprise the backbone of the Japanese archipelago. The western part of Japan is composed Geology mainly of Cretaceous to Paleogene felsic volcanic and plutonic rocks and accretionary complexes. The Tectonics eastern part of the country is covered extensively by Neogene sedimentary and volcanic rocks. During Accretion Metamorphism Plutonism coastlines of the Japanese Islands. These geological units are divided by age and origin: i.e. Paleozoic con- Back arc tinental margin; Paleozoic island arc; Paleozoic accretionary complexes; Mesozoic to Paleogene accre- Collision tionary complexes and Cenozoic island arcs. These are further subdivided into the following tectonic units, e.g. Hida; Oki; Unazuki; Hida Gaien; Higo; Hitachi; Kurosegawa; South Kitakami; Nagato-Renge; Nedamo; Akiyoshi; Ultra-Tamba; Suo; Maizuru; Mino-Tamba; Chichibu; Chizu; Ryoke; Sanbagawa and Shimanto belts. The geological history of Japan commenced with the breakup of the Rodinia super continent, at about 750 Ma. At about 500 Ma, the Paleo-Pacific oceanic plate began to be subducted beneath the continental margin of the South China Block. Since then, Proto-Japan has been located on the convergent margin of East Asia for about 5oo Ma. In this tectonic setting, the most significant tectonic events recorded in the geology of Japan are subduction-accretion, paired metamorphism, arc volcanism, back-arc spreading and arc-arc collision. The major accretionary complexes in the Japanese Islands are of Permian, Jurassic and Cretaceous-Paleogene age. These accretionary complexes became altered locally to low-temperature and high-pressure metamorphic, or high-temperature and low-pressure metamorphic rocks. Medium- pressure metamorphic rocks are limited to the Unazuki and Higo belts. Major plutonism occurred in Paleozoic, Mesozoic and Cenozoic time. Early Paleozoic Cambrian igneous activity is recorded as granites in the South Kitakami Belt. Late Paleozoic igneous activity is recognized in the Hida Belt. During Creta- the Takidani Granite intruded at about 1-2 Ma. During Cenozoic time, the most important geologic events are back-arc opening and arc-arc collision. The major back-arc basins are the Sea of Japan and the Shikoku and Chishima basins. Arc-arc collision occurred between the Honshu and Izu-Bonin arcs, and the Honshu and Chishima arcs. 2012 Elsevier Ltd. All rights reserved. 1. Introduction This new geological map is epoch-making, and will be the most use- ful map for geoscientists and other users who wish to understand the The international geological map of Asia at the scale of geology and tectonics of Asia. This special issue is designed to help 1:5,000,000 (IGM5000) will be completed in 2012. The map is the users understand easily the geology and tectonics of the Eurasian re- output of a project of the Commission for the Geological Map of gion. The author of this article is the compiler of the geological map the World (CGMW), organized by Ren Jishun of the Chinese Academy of Japan, and the project leader for Region V, which covers Japan, the of Science, Institute of Geology. This geological map covers most of Philippines, Malaysia, Indonesia and the offshore areas of East and the Eurasian continent from Japan in the east, to Turkey in the west. Southeast Asia. Hence, he has a responsibility to explain the geology of Japan as presented in this paper. The geological units in the Japanese Islands were reviewed by Iso- E-mail address: koji-wakita@aist.go.jp zaki et al. (2010a,b), while the tectonic history of the country in the 1367-9120/$ -see front matter ? 2012 Elsevier Ltd.All rights reserved. http://dx.doi.org/10.1016/jseaes.2012.04.014 76 K.Wakita/Journal of Asian Earth Sciences 72(2013)75-87 Phanerozoic era was described by Isozaki et al. (2011) and Maruyama ments, sedimentary, igneous and metamorphic rocks. Ages of major et al. (2011). Maruyama et al. (1997) have illustrated the paleo-geo- geological units represented range from Cambrian to Quaternary. graphic changes since 750 Ma. The author recently compiled a digital The clearly identified Precambrian rocks are gneiss of the clasts in geological map of Japan at a scale of 1:200,000, summarizing the coun- the Kamiaso conglomerate in the Jurassic accretionary complex of try's regional geology, based on the results of recent research (Wakita the Mino Belt (Adachi, 1971; Shibata and Adachi, 1974). The exis- et al., 2009). It has been made accessible online as a digital geological tence of a Precambrian basement is inferred, as the provenance of map (http://riodb02.ibase.aist.go.jp/db084/). In this paper, the author the sedimentary and metamorphic rocks of the Hida and Mino belts. reviews the geology and geological history of Japan, based on these re- The metamorphic rocks of the Higo and Hitachi belts intruded by cent geological documents, which have provided a valuable input to 500 Ma granites (Tagiri et al., 2011)are possible candidate ofthe Pre- the International Geological Map of Asia (IGMA500). cambrian basement. Recently, the Cambrian sedimentary rocks Were found in the Hitachi Belt(Tagiri et al., 2010, 2011). On the other 2. Ages of the major geological units hand, the youngest sediments and volcanic rocks can be found all over the country in alluvial plains and as active volcanoes. Although The geology of Japanese Islands is extremely complicated (Fig. 1). rocks of Paleozoic age are not widely distributed, rocks and forma- The rock types distinguished on the map are unconsolidated sedi- tions of late Mesozoic to Cenozoic can be found easily all over Japan 130E 140E 140°E Sapporo 40N Sea of Japan PacificPlate 40N Philipine Sea Plate 130°E Sendai SedimentaryMetamorphic Plutonid Koch Quaternary 1 13 Neogene 14 Paleogene 3 m： 388 15 Cretaceous 囍 6 16 18 Jurassic 10 8 19 Triassic 11 Permian 12 data.1: Quaternary sediments, 2: Neogene sedimentary rocks, 3: Paleogene sedimentary rocks and accretionary complex,4: Cretaceous sedimentary rocks and accretionary complex, 5: Jurassic sedimentary rocks and accretionary complex, 6: Triassic sedimentary rocks, 7: Permian sedimentary rocks and accretionary complex, 8: Paleogene Jurassic plutonic rocks, 20: ultramafic rocks. K.Wakita/Journalof AsianEarthSciences 72(2013)75-87 (Fig. 1). Rocks of Jurassic age occur mainly in Jurassic accretionary Some parts of this belt were subjected to medium-pressure complexes which comprise the backbone of the Japanese archipel- metamorphism between 250 and 240 Ma, caused by continent- ago. The western part of Japan is composed mainly of Cretaceous continent collision. These rocks were intruded by granitic rocks to Paleogene felsic volcanic and plutonic rocks and accretionary (Funatsu granite) at 220-180 Ma, and were covered by Jurassic to complexes. The eastern part is extensively covered by Neogene sed- Early Cretaceous shallow marine to non-marine formations. imentary and volcanic rocks. During the Quaternary, volcanoes erupted in various parts of Japan and alluvial plains formed all 2.1.2. Oki Belt (OK) around the coastlines of the Japanese Islands. a so de go e q pnns psi u si o These geological units are classified by age and origin of forma- the western extension of the Hida Belt, which extends to the north- tion: i.e. Paleozoic continental margin; Paleozoic island arc; Paleo- ern margin of the Chugoku and Kyushu regions. The belt is com- zoic accretionary complexes; Mesozoic to Paleogene accretionary complexes; and Cenozoic island arcs (Fig. 2). These are further sub- pressure metamorphic rocks and granites of Paleozoic age. The divided into traditional tectonic divisions. Stratigraphic columns metamorphic rocks are gneisses and migmatites of granulite to for these divisions are shown in Fig. 3. upper amphibolite facies. The protoliths of the metamorphic rocks were granite, basalt, sandstone, mudstone and limestone. The ages 2.1. Paleozoic continental margin of formation of the protoliths were Carboniferous to Permian. The age of metamorphism is about 250 Ma, CHIME from monazite (Su- 2.1.1. Hida Belt (HD) zuki and Adachi, 1994). Zircon cores from gneiss yielded 300- The shallow marine to non-marine sedimentary rocks, such as 350 Ma ages, while the Sm-Nd isochron age of amphibolite is limestone, sandstone and mudstone of this belt were deposited 2000 Ma (Tanaka and Hoshino, 1987). during Middle to Late Paleozoic time along the Asian continental margin. These Paleozoic formations were intruded by granites at 2.1.3. Unazuki Belt (UZ) 330-300 Ma and 270-250 Ma, and were subjected to high-temper- This belt is located along the eastern margin of the Hida Belt, ature and low-pressure metamorphism (Kunugiza et al., 2000). and is characterized by medium-pressure metamorphic rocks such Cenozoic island arc Cretaceous to Paleogene accretionary complex Jurassic accretionary complex Middle Paleozoic accretionary complex Late Paleozoic accretionary complex Paleozoic arc Northeast Japan uede Continental fragment and ophiolite OK Fig. 2. Major tectonic units of Japan (modified from Isozaki (2010), Figs. 1 and 4). AK: Akiyoshi Belt, CB: Chichibu Belt, CS: Chishima Arc, CZ: Chizu Belt, HD: Hida Belt, HG: OE: Oeyama Ophiolite, OK: Oki Belt, RK: Ryoke Belt, SB: Sanbagawa Belt (SB1: mid-Cretaceous; SB2: Late Cretaceous), SK: South Kitakami Belt, SM: Shimanto Belt, SO: Suo Belt, UN: Unazuki Belt, UT: Ultra-Tamba Belt. 78 K.Wakita/JournalofAsianEarthSciences72(2013)75-87 as staurolite schist (Hiroi, 1983). U-Th-Pb-EMP chemical age of metamorphism of the same age and type as those of the Hida Belt. uraninite is 240 Ma. The rocks of the belt were subjected to the Hence, this belt is sometimes regarded as part of the Hida Belt Hida Maizuru S. Kitakami Kurosegawa (Ma) Quaternary 2.6 Neogene nonmarineformations 23.0 harineformations Paleogene 65.5 ophiolite volcanism Jurassic plutonism 199.6- Triassic metamorphism 251.0 Permian 299.0 Carboniferous 359.2 Devonian 416.0 Silurian Ordovician Cambrian B Belt Akiyoshi Mino-Tamba Chichibu Shimanto Chishima Age Quaternary 2.6 Neogene 23.0 Paleogene 65.5 Cretaced Jurassic nonmarineformations 199.6 Triassic marineformations 251.0 Permian 299.0 Carboniferous 359.2 Devonian olutonism Silurian Ordovician metamorphism Cambrian Fig. 3. Tectono stratigraphic columns of the major tectonic units. (A) Hida, Maizuru, South Kitakami, Kurosegawa belts. (B) Akiyoshi, Mino-Tamba, Chichibu and Shimanto belts, and Chishima Arc. K.Wakita/Journalof Asian EarthSciences 72(2013)75-87 (Kunigiza et al., 2006). The protoliths of the metamorphic rocks (2010) respectively. U-Pb zircon ages from metasomatic jadeitite were sandstone, mudstone, limestone and basalt, formed around is about 520 Ma (Kunugiza and Goto, 2010). 300 Ma on the ancient rifted continental shelf. Carboniferous bry- Ozoans were found in limestone of this belt (Hiroi et al., 1978). Iso- 2.2.3. Kurosegawa Belt (KS) zaki et al. (2010b) regard the Unazuki Belt as an extension of the continent-continent collision zone between the North and South This belt is located in the Outer Zone of Southwest Japan, and is China blocks and the Higo and Hitachi belts. closely associated with the Jurassic accretionary complex of the Chichibu Belt. It is regarded as serpentinous mélange, including 2.1.4. Higo Belt (Hig) various types of rocks, such as low temperature high-pressure metamorphic rocks, granites, as well as Paleozoic and Mesozoic The Higo Belt, located in Kyushu Island in southwest Japan, con- sists mainly of metamorphic rocks, and Paleozoic to Mesozoic sed- formations. The main components of this belt are the Terano low-temperature and high-pressure metamorphic rocks (400 Ma), imentary rocks. The metamorphic rocks were originally medium- the Mitaki Granite (about 440 Ma) and Silurian to Devonian, Car- pressure metamorphic rocks with isotopic ages of about 230 Ma. They were overprinted by high-temperature and low-pressure boniferous to Permian and Triassic to Cretaceous formations. The Paleozoic to Mesozoic formations are correlated with some compo- metamorphism (Ryoke Metamorphism) in the Cretaceous. Re- cently, a very old granite (about 500 Ma) was reported from the nents of the South Kitakami Belt. The low-temperature and high- Higo Belt (Sakashima et al., 1999). The sedimentary rocks are mar- pressure metamorphic rocks are the same as those found in the Nagato-Renge Belt. The sepentinous matrix was originally derived ine deposits of Permian and Cretaceous in age. The rocks of the Hit- achi Belt of northeast Japan, as well as the rocks of the Unazuki Belt from the Oeyama Ophiolite. are regarded as the equivalent of the this belt. 2.2.4. Oeyama Ophiolite (OE) 2.1.5. Hitachi Belt (Hit) Oeyama is the oldest ophiolite in Japan. It is distributed along Hitachi Belt is composed of weakly metamorphosed sedimen- the northern margin of the Maizuru Belt to the north of Kyoto tary, volcanic and plutonic rocks. The metamorphic grade ranges (Kurokawa, 1985; Ishiwatari, 1989). Isozaki and Maruyama from greenschist to amphibolite facies. The sedimentary rocks con- (1991) regarded the ophiolite as an isolated unit in Japan. How- tain Early Carboniferous corals and Early Permian fusulinids. Saka- ever, most Japanese researchers include the ophiolite in the Naga- shima et al. (2003) reported 491 Ma SHRIMP age for a sheared to-Renge Belt. In order to describe the island arc components and granite which was intruded into the surrounding metamorphic accretionary complexes independently, the Oeyama ophiolite is rocks. Tagiri et al. (2010, 2011) also reported several SHRIMP zir- considered to be a separate tectonic unit in this paper. The ophio- con ages such as 510 Ma for tuffaceous schist, 507 Ma for meta- lite is composed mainly of peridotite, dunite, harzburgite and andesite, 505 Ma for meta-porphyry which is intruded into meta- hornblendite with metagabbro and diabase. It does not have clear volcanic rocks of the Hitachi Belt. layering, but preserves a complete section of the original ophiolite sequence. The ophiolite is partly altered to serpentinite, containing 2.2. Paleozoic island arc system blocks of low-temperature and high-pressure metamorphic rocks, derived from the Nagato-Renge Belt (Tsujimori, 1999; Tsujimori 2.2.1. South Kitakami Belt (SK) and Itaya, 1999; Ishiwatari and Tsujimori, 2003). K-Ar dating The South Kitakami Belt was originally defined by Shimazu yields 469-444 Ma ages (Nishimura and Shibata, 1989). The ser- et al. (1970) and later by Ehiro and Suzuki (2003) and Ehiro et al. pentinite of the Kurosegawa Belt is correlated with this ophiolite (2005). This belt is located in the northeastern part of Japan. It is (Ishiwatari, 1985; Isozaki and Maruyama, 1991). characterized by Ordovician ultramafic to mafic rocks and shallow marine continental shelf deposits, ranging in age from Silurian to Jurassic. The basement is composed of the Hayachine Complex 2.2.5. Maizuru Belt (MZ) (421-484 Ma), Tsubonosawa metamorphic rocks (334 Ma, Rb-Sr This belt is located in the Inner Zone of the Southwest Japan. It whole rock age), Hikami granite (412 Ma, Rb-Sr whole rock age), is composed of the Yakuno mafic rocks, Yakuno felsic rocks and Nishidouhira metamorphic rocks (511 Ma, U-Pb zircon age) and Permian to Triassic sedimentary rocks (Ishiwatari, 1985; Hayasaka, Cambrian granite (520 Ma). The Paleozoic formations are com- 1990. The Yakuno mafic rocks are called the Yakuno Ophiolite. posed of limestone, mudstone, sandstone, conglomerate, basalt They are composed of mudstone, metabasalt, amphibolite, meta- and tuff. The limestone sometimes yields corals and fusulinids. gabbro, metagranite and ultramafic rocks (Ishiwatari, 1985; Ish- The Mesozoic formations are mainly shallow marine deposits, iwatari and Tsujimori, 2003). The sheeted basalt dykes, pillow yielding ammonoids and bivalves. lavas, and hyaloclastites in this belt are interbedded with mud- stone. Metagabbro was metamorphosed in the amphibolite-gran- 2.2.2. Hida Gaien Belt (HG) ulite facies to the pyroxene-granulite facies. Using K-Ar dating, This belt is distributed along the southeastern margin of the the age of the amphibolite was determined as 280-250 Ma or Early Hida Belt (Kamei, 1955; Tsukada, 2003; Tsukada et al., 2004). It to Middle Permian. Two types of basalt are recognized in this belt is also called the Hida Marginal Belt. It is a mixture of various rocks such as island arc basalt and ocean-ridge basalt. These rocks and formations, which are correlated with the South Kitakami, originated in a Permian island arc and back-arc basin (Ishiwatari, Akiyoshi, Maizuru and Nagato-Renge belts and the Oeyama Ophi- 1999). olite Kumugiza et al., 2004; Kunugiza and Goto, 2010; Kunugiza Yakuno felsic rocks are made of granite, granodiorite, quartz and Maruyama, 2011). The components are Ordovician to Permian diorite, tonalite, metagabbro, dolerite and amphibolite. The Rb-Sr continental shelf sediments of the South Kitakami Belt, Permian total isochron ages ranges from 425 to 149 Ma, and the U-Pb SHRIMP ages from zircon are 249 and 243 Ma. The lower part of nic rocks of the Maizuru Belt, ultramafic rocks of Oeyama Ophiolite the Permian formation is composed of basalt with mudstone, while and Jurassic to Cretaceous shallow marine to non-marine sedi- the middle and upper parts are mainly mudstone, associated with ments, which also cover the Hida Belt. Ordovician conodonts, and sandstone and limestone. The Triassic formation is composed of U-Pb age of igneous zircon of latest Early Ordovician (472 Ma) conglomerate and sandstone with mudstone, and yields bivalves were reported by Tsukada and Koike (1997) and Nakama et al. and ammonites. 80 K. Wakita/Journal of Asian Earth Sciences 72 (2013) 75-87 2.3. Paleozoic accretionary complexes 2.3.5. Suo Belt (SO) This belt is located in the Inner Zone of Southwestern Japan. It 2.3.1. Nedamo Belt (ND) forms the main part of the Sangun Metamorphic Belt (Nishimura This belt is located along the northwestern margin of the South et al., 1989; Nishimura, 1990, 1998), which has been divided into Kitakami Belt (Ehiro and Suzuki, 2003; Uchino et al., 2005） in three belts. These are the Nagato-Renge, Suo and Chizu belts, based Northeast Japan. This is the oldest of the non-metamorphosed on the ages of the metamorphic events (Nakajima et al., 1992). The accretionary complexes in Japan, formed during Early Carbonifer- Suo Belt is composed mainly of low-temperature and high-pres- ous accretion. It is composed of sandstone, mudstone, felsic tuff sure metamorphic rocks, formed at about 200 Ma, associated with and basalt. It is characterized by the dominant presence of felsic metagabbro. This belt consists of pelitic, siliceous and mafic crys- volcanic rocks, unlike the other accretionary complexes in Japan. talline schists. The protoliths of these metamorphic rocks formed The sandstone contains detrital grains of volcanic arc origin, and the Permian accretionary complex of the Akiyoshi Belt. Metamor- the conglomerate yields pebbles of low-temperature and high- phic grades are mainly the pumpellyite-actinolite and glauco- pressure metamorphic (347-317 Ar-Ar age) and ultramafic rocks. phane schist facies, and partly the epidote-amphibolite facies. These grains and pebbles were derived from the South Kitakami The metamorphic ages are 227-191 Ma (K-Ar age from muscovite) Belt (Uchino et al., 2008). This accretionary complex is the proto- (Nishimura et al., 1989; Shibata and Nishimura, 1989, and 253- lith of the metamorphic rocks of the Nagato-Renge Belt (Uchino 237 Ma (K-Sr age from hornblende in metagabbro (Shibata et al., et al., 2005; Kawamura et al., 2007). 1977; Nishimura and Shibata, 1989). 2.3.2. Nagato-Renge Belt (NR) 2.4.Mesozoic accretionary complexes This belt is located at the margin of the Hida Belt in the western and central regions Japan. It is composed of low-temperature and 2.4.1. Mino-Tamba Belt (MT) high-pressure metamorphic rocks of the glaucophane schist facies, This belt, distributing widely in Japan, is one of the major com- the pumpellyite-actinolite facies and the lawsonite-pumpellyite ponents of the Japanese Islands (Figs. 1 and 2). It is a Jurassic accre- facies (Tsujimori, 2002). The protoliths of the metamorphic rocks tionary complex, but the age of accretion ranges from Late Triassic were mostly mafic igneous rocks, with some felsic igneous and to earliest Cretaceous. It consists of sandstone, mudstone, con- ultramafic rocks. The metamorphic ages of the rocks are 298- glomerate, siliceous shale, chert, limestone and basalt. Limestone 279 Ma (K-Ar) and 327-273 Ma Rb-Sr (mostly 320 Ma) age from associated with basalt yields Carboniferous Permian microfossils phengite and 264-303 Ma (K-Ar) from muscovite. The protoliths such as fusulinids and corals, while chert includes Carboniferous of the metamorphic rocks of this belt were Paleozoic accretionary to Triassic conodonts and up to Early Jurassic radiolarians. Siliceous complexes, which were contemporaneous with the Nedamo Belt shale and mudstone contains Triassic to Early Cretacous radiolari- (Uchino et al., 2005; Isozaki et al., 2010a). The northeastern exten- ans which indicates the duration of sedimentary accretion at the sion of this belt is traced along the South Kitakami Belt into the ancient trench. These rocks were components of Ocean Plate Stra- Matsugadaira, Sanjo and Tateishi metamorphic rocks (380 Ma) tigraphy (Matsuda and Isozaki, 1991; Wakita, 2000a; Wakita and (Uchino et al., 2008). Metcalfe, 2005). The sedimentary sequence is tectonically dis- rupted and imbricated, and sometimes chaotically mixed, to form the complex structures of an ancient accretionary prism. Jurassic 2.3.3. Akiyoshi Belt (AK) accretionary complexes of similar components and structures also This belt is distributed in several isolated areas of the Inner form the Chichibu Belt, and their metamorphic equivalents occur Zone of Southwestern Japan, extending into the western margin in the Chizu, Ryoke and Sanbagawa belts. of the Hida Belt. It forms a Permian accretionary complex, com- posed of sandstone, mudstone, conglomerate, siliceous shale, felsic tuff, chert, limestone and basalt. The Akiyoshi Limestone (Toriy- 2.4.2. Chichibu Belt (CB) ama, 1954), a famous component of this belt, yields well-preserved The Chichibu Belt forms a major component of the Outer Zone fossils (e.g. fusulinids and corals) ranging from Tournaisian to of Southwestern Japan. It is a Jurassic accretionary complex, like Capitanian in age (Toriyama, 1958; Haikawa and Ota, 1978), and the Mino-Tamba Belt (Matsuoka, 1992). The age of accretion is regarded having formed as a reefal limestone on a seamount in ranges from Early Jurassic to Early Cretaceous. The youngest part the Panthalassan Ocean (Kanmera and Nishi, 1983; Igawa, 2003). is sometimes called the Sambosan Belt, which is characterized by This limestone is an allochthonous block, which was mixed with mélanges, including blocks of Permian and Triassic limestone. trench fill sediments during Permian accretionary processes (Sano The Jurassic accretionary complex is composed of sandstone, mud- and Kanmera, 1988, 1991a,b). Major accretion occurred during the stone, conglomerate, siliceous shale, chert, limestone and basalt. Middle to Late Permian. Many components of this belt are unmeta- The tectonic stacking of Ocean Plate Stratigraphy is well-exposed and well-documented in the Togano Group (Matsuoka, 1984) of metamorphism of the prehnite-pumpellyite and pumpellyite- the Shikoku region. The North Kitakami Belt is equivalent to this actinolite facies, the same as the metamorphism of the Suo Belt. belt. 2.3.4. Ultra-Tamba Belt (UT) 2.4.3. Chizu Belt (CZ) This belt is located between the Maizuru Belt and Tamba (Mino- This belt is distributed in the Inner Zone of the Southwest Japan. Tamba) Belt (Caridroit et al., 1985). This is also a Middle to Late It was called the Sangun Metamorphic Rocks in the past. However, Permian accretionary complex, similar to the Akiyoshi Belt. How- based on the ages of the protolith and the metamorphism, the Chi- ever, the tectonic position and some of the components are differ- zu Belt was defined by Shibata and Nishimura (1989) and further ent. The Ultra-Tamba Belt is composed of Late Carboniferous to by Isozaki et al. (2010a). It comprises low-temperature and high- Early Permian basalt and conglomerate, sandstone, mudstone, pressure metamorphic rocks of the greenschist and glaucophane felsic tuff and chert of Middle to Late Permian in age. The Middle schist facies. The metamorphic age is about 180 Ma. The protoliths Triassic clastic Shimamoto Formation is partly intercalated in this of these metamorphic rocks formed part of the Jurassic accretion- belt. ary complex of the Mino-Tamba Belt (Hayasaka, 1987). K. Wakita/Journal of Asian Earth Sciences 72 (2013) 75-87 81 2.4.4. Ryoke Belt (RK) Trench. Huge amounts of rock previously accreted were removed This belt is a famous regional metamorphic belt as part of a by tectonic erosion (Suzuki et al., 2010; Yanai et al., 2010). On paired metamorphic belt (Miyashiro, 1961), together with the San- the other hand, the young Philippine Sea Plate is being subducted bagawa belt. It comprises the Ryoke metamorphic rocks, which and forming an accretionary wedge along the Nankai Trough in Southwestern Japan. The collision between the Honshu and Izu phism at about 100 Ma (Cretaceous) and felsic plutonic rocks (Iso- arcs provided trench-fill sediments for the accretionary wedges zaki and Maruyama, 1991). The Ryoke Belt was located at the and compression which uplifted the accretionary wedges to form eastern end of the Eurasian continental margin in Cretaceous times the Boso Peninsula. Extensive volcanic activity has occurred due (Nakajima, 1994, 1996). to plate subduction beneath the northeastern Honshu Arc. The The highest grade of metamorphism reached was the granulite strong, 9.0 magnitude, earthquake in northeast Japan in March facies. The protoliths of the metamorphic rocks formed the Jurassic 11, 2011 was triggered by subduction of the Pacific Plate beneath accretionary complex of the Mino-Tamba Belt. The lowest grade the Honshu Arc along the Japan Trench. Reverse fault movement part of the Ryoke Belt passes gradually into rocks of the Mino-Tam- between the Honshu Arc and the Pacific Plate was the main cause of the massive tsunamis. shi Belt was also subjected to the Ryoke metamorphism. The felsic In the southwestern part of the Honshu Arc, volcanic activity is plutonic rocks are granite, granodiorite and tonalite with 100- not as active, as in the north east, except in the Kyushu region, 85 Ma CHIME ages (Suzuki and Adachi, 1998; Suzuki et al., where extension tectonics causes rifting and active volcanism. 1999). Most of these granitic rocks are of I-type but some of them Sometimes earthquakes occur along the Nankai Trough as the re- are of S-type. sult of plate subduction. 2.4.5. Sanbagawa Belt (SB) (older sub-belt = SB1, younger sub- 2.5.2. Ryukyu Arc (RK) belt=SB2) The Ryuku Arc is underlain by Paleozoic and Mesozoic accretion- The Sanbagawa Belt is formed of low-temperature and high- ary complexes, belonging to the Nagato-Renge, Chichibu and Shim- pressure metamorphic rocks. Isozaki et al. (2010b) divided the anto belts. The Okinawa Trough, a back-arc basin, lies along the Sanbagawa Belt into two sub-belts, based on the ages of the proto- northwestern side of the Ryukyu Arc. This arc extends southwest- lith and metamorphism. The protolith of the older sub-belt was the ward as far as Taiwan, where it connects with the Philippine Arc. earliest Cretaceous accretionary complex of the Chichibu Belt, and 2.5.3. Izu-Bonin Arc (IB) morphism of the greenschist to eclogite facies during mid-Creta- The Izu-Bonin volcanic arc results from the subduction of the ceous time (Aoki et al., 2008). On the other hand, the protolith of Pacific Plate beneath the Philippine Plate. It is sometimes called the metamorphic rocks of the younger sub-belt was the mid-Creta- the Izu-Bonin-Mariana Arc, extending for about 1200 km long, ceous accretionary complex of the Shimanto Belt. This belt was with a width of about 400 km. The thickness of the arc is over subjected to low-temperature and high-pressure metamorphism 20 km (Takahashi et al., 1998). It forms the eastern half of the Pa- of the greenschist to the eclogite facies during Late Cretaceous time leo-Izu-Bonin Arc. Back-arc spreading caused the Paleo-Izu-Bonin (60-70 Ma). The metamorphic rocks of the Kamuikotan and Tokoro Arc to divide into the present Izu-Bonin Arc and the Kyushu-Palau belts of Hokkaido can be regarded as contemporaneous with the Ridge, and to form the Philippine Sea Plate between the Paleocene Sanbagawa Belt. The Abukuma Belt of Northeast Japan is the equiv- and the Miocene. Seismic imaging suggests that most of the pres- alent of this belt. ent Izu-Bonin Arc crust was created during Eocene-Oligocene time (Kodaira et al., 2008, 2010). A Paleogene-Neogene volcanic arc 2.4.6. Shimanto Belt (SM) forms the basement of the Izu-Bonin Arc, overlain by Quaternary The Shimanto Belt is a mid-Cretaceous to Miocene accretionary volcanic and sedimentary rocks. complex in the Outer Zone of Southwestern Japan (Taira et al., 1988). It is subdivided into a mid- to Late Cretaceous northern belt 2.5.4. Chishima Arc (Kurile Arc) (CS) and a Paleogene to Miocene southern belt. The southern belt is sometimes called the Setogawa Belt. The Cretaceous to Paleogene The Chisima Volcanic Arc extends from the Eastern part of Hok- kaido to the Aleutian Islands via the Kurile Islands. It caused by the accretionary complex of this belt is composed mainly of tectoni- subduction of the Pacific Plate beneath the Okhotsk Plate. The cally stacked turbidites, associated with tectonic melanges includ- ing blocks of chert, limestone and basalt in a pelitic matrix. basement of the arc is formed of Cretaceous igneous and sedimen- tary rocks, some of which are exposed in eastern Hokkaido. Ceno- Northern extensions of the belt are the Hidaka, Kamuikotan and zoic volcanic and sedimentary rocks cover the Cretaceous Tokoro belts in Hokkaido. basement extensively. This arc forms part of one of the most active 2.5. Cenozoic island arc system convergent margins in the world, with frequent volcanic eruptions and earthquakes. This arc became connected to the Honshu Arc in 2.5.1. Honshu Arc (HS) Hokkaido by the Neogene collision between the two arcs. After the opening of the Sea of Japan, the country became an is- land arc isolated from the Asian continent. The main parts of the 3. Tectonic events in Japan during the Phanerozoic Japanese Islands are called the Honshu Arc, including not only Honshu, but also Kyushu, Shikoku and Hokkaido islands. The base- 3.1.Outline ment of the arc is composed of Paleozoic to Cenozoic rocks ac- creted by ocean plate subduction along the Asian continental The geological history of Japan commenced with the breakup of margin. The Honshu Arc is divided into the Southwest Honshu the supercontinent of Rodinia at about 750 Ma (Isozaki et al., (or Japan) Arc and the Northeast Honshu (or Japan) Arc by the Fos- 2010a). The South China and North American blocks separated, sa Magna in central Japan. and the Paleo-Pacific Ocean was formed between them. Japan The Pacific Plate is being subducting along the Japan Trench be- was developed on the passive margin of the South China Block. neath the northeastern part of the Honshu Arc. Subduction of the About 500 Ma (Isozaki et al., 2010a), The Paleo-Pacific Oceanic s on ps ed onno p pe Plate was subducted beneath the continental margin of the South 82 K.Wakita/Journal of Asian Earth Sciences72(2013)75-87 accretion, tectonic erosion, igneous activity have occurred in asso- for about 500 million years. This means that the most significant tectonic events recorded in the geology of Japan are subduction- of Japan. However, the geological record is deficient in the Early to accretion tectonics, as well as tectonic erosion. The earliest geolog- Middle Paleozoic. There is more geological information on sedi- ical record in Japan is the formation of ophiolitic rocks at about ment accretion during the Late Paleozoic and the Mesozoic to 500 Ma, associated with low-temperature and high-pressure meta- Paleogene. The distribution and a stratigraphic column to illustrate morphic rocks. These are remnants of a volcanic arc that was the protoliths of these accretionary complexes are shown in Fig. 5. formed near the Asian continental margin. In this section, some These protoliths are composed of similar lithologies and strati- of the major tectonic activities that occurred throughout the geo- graphic sequences but of different ages. Components of the accre- logical history of Japan are described, such as sediment accretion, tionary complexes are derived from Ocean Plate Stratigraphy metamorphism, plutonism, back-arc spreading and arc-arc colli- (Matsuda and Isozaki, 1991; Wakita and Metcalfe, 2005). Ocean sion (Fig. 4). For tectonic erosion, which is also one of the most Plate Stratigraphy is composed of rocks of seamounts, pelagic important events in Phanerozoic Japan, refer to several recent pub- and hemipelagic sediments of the ocean floor and trench turbi- lications (e.g. Suzuki et al., 2010). dites. Basalt and limestone are derived from the upper part of sub- ducting seamounts. Radiolarian cherts are pelagic sediments 3.2.Accretionary process deposited on the seafloor, and siliceous shale with felsic tuff was deposited in a hemipelagic environment. Sandstone, mudstone Japan or Proto-Japan has been located on a convergent margin and conglomerate are trench fill sediments, sourced as detrital since about 5o0 Ma. Various tectonic processes such as sediment grains from the continental or island arc side. All these rocks: ba- salt, limestone, chert, siliceous shale, felsic tuff, sandstone, mud- stone and conglomerate, form the basic components of the accretionary complexes. These components are detached from the oceanic plate and tectonically stacked to form an accretionary Accretion Metamorphism Plutonism wedge near the trench, where they are dismembered, fragmented, and mixed together to form complex structures in the accretionary wedges. Quaternary The oldest non-metamorphosed accretionary complex is of Car- offshore boniferous age in the Nedamo Belt. The protolith of the low-tem- Takidani (1.7Ma) perature high-pressure metamorphic rocks of the Nagato-Renge Belt were incorporated into accretionary complexes of Middle to Late Paleozoic age. Neogene Two of the major accretionary complexes are the Permian Outer Zone Akiyoshi and Ultra-Tamba belts. Sediment accretion in these belts South Shimanto San-in Belt Paleogene Shimanto North CretaceousAC Shimanto Ryoke Cretaceous Sanbagay Ryoke, San-yo Hida Mino - Tamba Chizu Permian Belt Jurassic AC Suo Hida CretaceousAC Triassic Hida Akiyoshi Nagoto Unazuki Permian -Renge = Carboniferous Hida sandstone Perr South Kitakami siliceous shale Devonian Kurosegawa chert limestone Jurassic AC Silurian ifero basalt Permian AC boni Ordovician Car Cambrian Fig. 4. Major tectonic events in Japan: accretion, metamorphism and plutonism Fig. 5. Major accretionary complexes of Japan and the stratigraphy of their (modified from Isozaki et al. (2011), Figs. 1 and 14). protoliths (after Wakita and Metcalfe, 2005, Fig. 6). K. Wakita/Journal of Asian Earth Sciences 72 (2013) 75-87 83 occurred during the Middle to Late Permian. The components of ultramafic rocks. Cenozoic accretionary complexes were also these accretionary wedges are Carboniferous to Middle Permian developed in the offshore area between the Nankai Trough and limestone associated with basalt, Permian chert including radiolar- the mainland of Japan. The structure of the accretionary wedges ian remains and sponge spicules, felsic tuff, sandstone, mudstone was verified using seismic surveys and ocean drilling. and conglomerate. These rock types were tectonically stacked to form the accretionary complex, together with the chaotic mélanges 3.3.Metamorphisim of the Tsunemori Formation. Sediment accretion was contempora- neous with the collision between the North and South China Metamorphic events in Japan are divided into low-temperature blocks. The collision was a the possible trigger for the provision and high-pressure and high-temperature and low-pressure types. of sediments to the ancient trench which now form the Permian Medium-pressure metamorphism is recognized only in limited accretionary wedges of the Akiyoshi and Ultra-Tamba belts. Accre- areas such as in Hida and Unazuki belts. As ocean plates have been tionary complexes were neither formed nor tectonically eroded subducted beneath Japan since 500 Ma, low-temperature and during Early Triassic time. high-pressure metamorphism has occurred throughout Phanero- Jurassic accretionary complexes are extensively developed in zoic time. Low-temperature and high-pressure metamorphism oc- Japan (Wakita, 1988; Otsuka, 1988; Isozaki et al., 1990; Isozaki, a sn s 1997; Nakae, 1993). Although the accretion occurred from Late Tri- developed (Fig. 5). Maruyama et al. (2010) explained that low-tem- assic to Early Cretaceous, they are termed “Jurassic accretionary perature and high-pressure metamorphism was caused by the sub- complexes" in this paper. Jurassic accretionary complexes occur duction of ocean ridges, when the boundary of two different ocean metamorphosed in the Ryoke, Chizu and Sanbagawa belts. Jurassic plates were subducting along the trench. accretionary complexes are characterized by the dominance of In Japan Paleozoic metamorphism is recorded in the Hayachine radiolarian cherts. Oceanic Plate Stratigraphy was separated from complex (421-484 Ma, the Nishidohira metamorphic rocks the underlying ocean floor by a decollement in the super-anoxic (511 Ma) in the South Kitakami Belt, the Terano metamorphic claystone at the Permian-Triassic boundary (Isozaki, 1993, 1984). rocks (400 Ma) the Kurosegawa Belt and the Nagato-Renge meta- Therefore, sediments younger than P-T boundary claystone were morphic rocks (about 320 Ma. During the Triassic (200 Ma), low- accreted to form the accretionary complexes, which contain radio- temperature and high-pressure metamorphism occurred in the larian chert and hemipelagic siliceous shale, as well as trench Suo and Akiyoshi belts. During the mid-Cretaceous (110- turbidites. These sediments were tectonically stacked by the offs- 120 Ma), the earliest Cretaceous accretionary complex of the San- craping process (Kimura and Hori, 1993; Isozaki et al., 1990). bagawa Belt was subjected to low-temperature and high-pressure Permian chert, limestone, and basalt were subducted and under- metamorphism in the greenschist to eclogite facies. During the plated in the deeper part of the accretionary wedge. This accretion- Late Cretaceous (60-70 Ma), the mid-Cretaceous accretionary ary complex is also developed in the Russian Far East and complex of the Shimanto Belt was subjected to low-temperature Northeast China (Kojima, 1989; Kojima et al., 2000; Yamakita and high-pressure metamorphism. This type of metamorphism oc- and Otoh, 1998; Otoh et al., 1999), and Palawan Island in the Phil- curred in the Sanbagawa Belt. ippines (Zamoras and Matsuoka, 2000). High-temperature and low-pressure metamorphism occurred Cretaceous to Paleogene accretion was the main cause of the in the Hida and Ryoke belts, caused by granitic intrusion between accretionary complex of the Shimanto Belt. Contemporaneous 330-300 Ma and 270-250 Ma (Kunugiza et al., 2000). The same accretionary wedges occur also in Indonesia (Wakita, 2000b). The metamorphism occurred in the Unazuki Belt where medium-pres- accretionary complex of the Shimanto belt is characterized by tec- a n m s n a tonically stacked turbidites. Melanges, including radiolarian chert, the mid-Cretaceous (100 Ma), high-temperature and low-pressure are less dominant than those containing turbidites. The Creta- metamorphism, the Ryoke metamorphism affected some parts of ceous-Paleogene accretionary complex of the Shimanto Belt is the Jurassic and Permian accretionary complexes. dominant in turbidite and contains less chert and limestone. This difference was caused by the position of the decollement, which 3.4. Igneous activity developed at the toe of accretionary wedge. In the present Barba- dos accretionary wedge the décollement was developed mainly Igneous activity occurred during the Paleozoic, Mesozoic and within the hemipelagic sediment, because the porosity of the sed- Cenozoic (Fig. 4). Early Paleozoic igneous activity is recorded as iments, including radiolarian remains, is very low, compared with Cambrian granite (520 Ma) and the Hikami Granite (412 Ma) of the other stratigraphic units (Shipley et al., 1997). In this case, sed- the South Kitakami Belt (Adachi et al., 1992; Suzuki and Adachi, iments younger than the hemipelagic sediments were accreted to 1993). Late Paleozoic igneous activity is identified in the Hida form the accretionary wedge, while sediments older than the metagranites of 330-300 Ma and 270-250 Ma in the Hida Belt. hemipelagic unit were not accreted, but subducted more deeply, Triassic to Jurassic (220-180 Ma) igneous activity is overprinted sometimes into the mantle. The accretionary complex of the Shim- on the Hida Belt. anto Belt was formed in the same way. Lower Cretaceous igneous activity was the precursor of exten- Recently, Isozaki et al. (2010a) demonstrated that the protoliths sive igneous activity in Late Cretaceous to Paleogene time. These of some parts of the Sanbagawa metamorphic rocks were derived are the Kwanmon Group in the Chugoku region, the Sasayama from the Cretaceous accretionary complex of the Shimanto Belt. Group in the Kinki region, the Harachiyama Formation in the Toho- ku region, and the old Ryoke Granites. The Kwanmon Group is pan using seismic reflection survey, and demonstrated that the composed of non-marine sediments, including andesitic volcanic Shimanto Belt is connected with the Sanbagawa Belt at depth. It and volcaniclastic rocks. The Sasayama Group is composed of appears that at least some of the Sanbagawa Belt is the metamor- andesitic lavas and pyroclastic rocks, as well as detrital sediments phosed equivalent of the Cretaceous accretionary complex of the of Early Cretaceous age. The Harachiyama Formation is composed Shimanto Belt. of calc-alkaline to alkaline volcanic rocks, olivine basalt, pyroxene Neogene accretionary complexes are exposed along the south- andesite, amphibolite andesite, dacite and rhyolite, ern margin of the Cretaceous-Paleogene accretionary complex of During Cretaceous to Paleogene time, extensive igneous activity the Shimanto Belt. These accretionary complexes were formed be- Occurred in Japan (Fig. 6). Major plutons are the Ryoke (100- tween 6 and 4 Ma. They include the Mineoka Ophiolite containing 70 Ma), Hiroshima (San-yo) (110-70 Ma), San-in (68-37 Ma), 84 K.Wakita/Journal of Asian Earth Sciences72(2013）75-87 130°E 140°E usvolcanicrocks 40N Kitakami Gr. Abukuma Gr. San-in Gr. Ryoke Gr. Hiroshima (San-yo) Gr. Fig. 6. Distribution of Late Cretaceous to Paleogene granites and felsic volcanic rocks in Japan. Abukuma (110-90 Ma）and Kitakami Granites (120-110 Ma). proto Izu-Bonin Arc into the Izu-Bonin Arc and Kyushu Palau Ridge. olivine and clinopyroxene. Near the bottom of the granodiorite and tonalite, associated with quartz diorite and peg- was rotated clockwise to form the Sea of Japan. The rotation started matite. The San-in Granites belong to the magnetite series and in 20 Ma, and stopped in 15 Ma. At 15 Ma, the opening of all three are subdivided into the Inbi intrusives (68-51 Ma) and the major marginal basins around Japan ceased. Namariyama intrusives (39-37 Ma). Late Cretaceous to Paleogene The Pacific Plate started to subduct beneath the Philippine Sea Fig. 8 : Sketches of trough- Plate about 50 million years ago, and produced the Izu-Bonin 8 was found, suggesting The 'detailed study of rhyo-dacitic and dacitic composition, and are associated with Middle Miocene (15 Ma) (Sugimura, 1972). The Izu-Bonin Arc is andesite and lake deposits. The youngest plutonism is recorded pushing the Honshu arc northward, together with the migration in central Japan; the K-Ar hornblende age of the Takidani Granite, of the Philippine Sea plate (Yamazaki et al., 2010; Tamura, 2011). the youngest granite in the world, ranges from 1.93 to 1.20 Ma Tamura et al. (2010) integrated new geochemical results with re- (Harayama, 1994). cent geophysical imaging of the arc and concluded that Miocene plutonic rocks in the Izu Collision Zone were derived from the Eo- 3.5. Opening of back-arc basin and arc-arc collision cene-Oligocene middle crust, which was partially melted, remobi- lized, and rejuvenated during the collision. In 4 Ma, the Tanzawa During Cenozoic time, the most important geologic events were mountains collided and were transferred to the Honshu Arc from back-arc opening and arc-arc collision. The major back-arc basins Izu-Bonin Arc (Takahashi et al., 1994, 1998). In the same way, are the Sea of Japan, and the Shikoku and Chishima (Kuril) basins. the Izu Peninsula collided and was accreted to the Honshu Arc at Collision occurred between the Honshu and Izu-Bonin (Izu-Ogasa- 1 Ma. wara) arcs, and between the Honshu and Chishima (Kurile) arcs. The collision between the Honshu Arc and the Chishima Arc Until 30 Ma, proto-Japan was located on the continental margin Observation and sampling were of the Asian continent. At about 30 Ma, continental rifting com- ura, 1986, 1994; Komatsu et al., 1983, 1989). It happened in two menced on the Asian continental margin. At 25 Ma the Sea of Japan stages (Kimura, 2002). The first stage of the collision was caused and the Chishima (Kuril) basins started to open as marginal seas. At by the westward movement of the Okhotsk Plate against the Eur- the same time, the Shikoku Basin started to open, dividing the asia Plate. Dextral oblique collision occurred between these two K. Wakita/Journal of Asian Earth Sciences 72 (2013) 75-87 85 plates in Paleogene time. The second stage occurred between 13 Earth Science, Hiroshima University, vol. 19, 1-39 (in Japanese with English and 8 Ma, when the Chishima Arc collided as forearc sliver with abstract). Hayasaka, Y., 1990. Maizuru terrane. In: Ichikawa, K., Mizutani, S., Hara, I., Hada, S., the western part of Hokkaido. The oblique subduction of the Pacific Yao, A. (Eds.), Pre-Cretaceous Terranes in Japan, vol. 224. IGCP Publication, plate against the Chishima Arc was the main cause of the uplift of Osaka, pp. 81-95. the Hidaka Mountains in central Hokkaido. In this collision the two Hiroi, Y., 1983.Progressive metamorphism of the Unazuki pelitic schists in the Hida terrane,central Japan.Contribution to Mineralogy and Petrology 82,334-350. arcs merged. Hiroi, Y., Fuji, N., Okimura, Y., 1978. New fossil discovery from the Hida sn hn e s Japan Academy 54, 268-271. 4. Summary Igawa, T., 2003. Microbial contribution to deposition of Upper Carboniferous and Lower Permian seamount-top carbonates, Akiyoshi, Japan. Facies 48, 61-78. Ishiwatari, A, 1985. Igneous petrogenesis of the Yakuno ophiolite (Japan) in the context of thediversity of ophiolites.ContributionstoMineralogy and Petrology 89, 155-167. 1. The major geologic entities of Japan are a Paleozoic conti- Ishiwatari, A., 1989. Ophiolites of Japan. Journal of Geography 98, 290-303 (in Japanese with English abstract). nental margin, a Paleozoic island arc, a Paleozoic accretion- Ishiwatari, A., 1999.Fragment of Paleozoic oceanic island arc crust in the Inner Zone ary complex, a Mesozoic accretionary complex and a of southwest Japan: the Kamigori metagabbro body, Hyogo Prefecture. Memoir Cenozoic island arc system. no j 2. The Paleozoic continental margin is represented by the abstract). Ishiwatari, A., Tsujimori, T., 2003.Paleozoic ophiolites and blueschists in Japan and Hida, Oki, Unazuki, Higo and Hitachi belts. On the other Russian Primorye in the tectonic framework of East Asia. The Island Arc 12, hand, Paleozoic island arcs form the South Kitakami Belt, 190-206. the Hida Gaien Belt, the Kurosegawa Belt, the Oeyama Isozaki, Y., 1984. Superanoxia across the Permo-Triassic boundary: record in Ophiolite and the Maizuru Belt. D.J. (Eds.), Global Environments and Resources, vol. 17. Canadian Society of 3. Accretionary complexes were developed during the Middle Petroleum Geologists, Memoir, pp. 805-812. to Late Paleozoic, Jurassic and from Cretaceous to Paleo- Isozaki, Y., 1993. Superanoxia across the P/T boundary and mass extinction. EOS (Transactions of American Geophysical Union) 74, 330. gene times. These complexes are composed of accreted Isozaki, Y., 1997. Jurassic accretion tectonics of Japan. The Island Arc 6, 25-51. sediments, derived from Ocean Plate Stratigraphy. Isozaki, Y., Aoki, K., Nakama, T., Yanai, S., 2010a. New insight into a subduction- 4. Accretion, metamorphism, .igneous activity, back-arc related orogen: a reappraisal of the geotectonic framework and evolution of the Japanese Islands.Gondwana Research 18, 82-105. spreading and arc-arc collision were the major tectonic events in the development of Japan. These are very impor- and new geotectonic subdivision of the Japanese Islands. Journal of Geography tant events on convergent margins, and resulted in keys to 100, 697-761 (in Japanese with English abstract). Isozaki, Y.,Maruyama, S.,Aoki, K.,Nakama,T., Miyashita,A.,Otoh, S.,2010b. understand the geology of Asia. Geotectonic subdivision of the Japanese Islands revisited: categorization and definition of elements and boundaries of pacific-type (Miyashiro-type) Orogen. Journal of Geography 119, 999-1053 (in Japanese with English abstract). Isozaki, Y., Maruyama, S., Fukuoka, F., 1990. Accreted oceanic materials in Japan. Acknowledgements Tectonophysics 181, 179-205. Isozaki, Y., Maruyama, S., Nakama, T., Yamamoto, S., Yanai, S., 2011. Growth and shrinkage of an active continental margin: updated geotectonic history of the The author expresses his gratitude to the Academician Ren Japanese Islands. Journal of Geography 120, 65-99 (in Japanese with English Jishun of the Chinese Academy of Science, for giving him the abstract). opportunity to work with IGMA500o. 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Abstracts with Program, Shimazu, M., Tanaka, K., Yoshida, T, 1970. Geology of Taro District, with Geological 1999 Annual Meeting of the Japanese Association of Mineralogists,Petrologists s ) ds o s i e d i and Economic Geologists, Sendai, Japan, p. B10 (in Japanese). English abstract). Kunugiza, K., Maruyama, S., 2011. Geotectonic evolution of the Hida Marginal Belt, Shipley, T.H., Ogawa, Y., Blum, P., Ashi, J., Brueckmann, W., Filice, F., Fisher, A.T., Central Japan: reconstruction of the oldest Pacific type orogeny of Japan.Journal Goldberg, D., Henry, P., Housen, B., Jurado, M.J, Kastner, M., Labaume, P., Lair, T., Leitch, E.C., Maltman, A.J., Meyer, A., Moore, G.F., Moore, J.C., Peacock, S., Kurokawa,K,1985.Petrology of the eyama ophiolitic complex nthe Inner zoneof Rabaute, A., Steiger, T.H., Tobin, H.J, Underwood, M.B., Su, Y, Yin, H., Zhen, Y., Zwart, G, 1997.Proceedings of the cean Drilling Program, scientific results, Mineralogy 6, 37-113. Northern Barbados Ridge; covering Leg 156 of the Cruises of the Drilling Vessel Maruyama, S., Isozaki, Y., Kimura, G., Terabayashi, M, 1997.Palegeographic maps JOIDESResolution,Bridgetown,Barbados,to Bridgetown,arbados,sites 947- of the Japanese Islands: plate tectonic synthesis from 750 Ma to the Present.The 949, 24 May-24 July 1994. Proceedings of the Ocean Drilling Program, Scientific Island Arc 6, 121-142. Results, vol. 156, 366 p. Maruyama, S., Masago, H., Katayama, I., Iwase, Y., Toriumi, M., Omori, S., Aoki, K., Sugimura, A., 1972. Plate boundaries in Japan and its adjacent areas. Kagaku 42, 2010. A new perspective on metamorphism and metamorphic belts. Gondwana 193-202 (in Japanese). Research 18, 106-137. Suzuki, K., Adachi, M., 1993. Th, U and Pb analytical data of monazites, zircons and Maruyama, S., Omori, S., Senshu, H., Kawai, K, Windley, B.F., 2011. Pacific-type Xenotimes from the Okuhinotsuchi granitic mass in the South Kitakami Terrane. Bulletin of the Nagoya University Furukawa Museum 9, 47-57. Journal of Geography 120, 115-223 (in Japanese with English abstract). Suzuki, K., Adachi, M., 1994. Middle Precambrian detrital monazite and zircon from Matsuda, T.,Isozaki, Y, 1991.Well-documentated travel history ofMesozoic elagic the Hida gneiss on Oki-Dogo Island, Japan: their origin and implications for the cherts in Japan: from remote ocean to subduction zone. Tectonics 10, 475-499. correlations of basement gneiss of Southwest Japan and Korea.Tectonophysics 235,277-292. part of Kochi Prefecture,Southwest Japan.Journal of Geological Society of Japan Suzuki, K., Adachi, M., 1998. Denudation history of the high T/P Ryoke metamorphic 90, 455-477 (in Japanese with English abstract). Matsuoka, A., 1992. jurassic and Lower Cretaceous tectonic evolution of the granitoids. Journal of Metamorprhic Geology 16, 23-37. Southern Chichibu terrane, southwest Japan. Palaeogeography, Suzuki, K, Adachi, M., Kato, T., Yogo, S., 1999. CHIME dating method and itS Palaeoclimatology, Palaeoecology 96, 71-88. application to the analysis of evolutional history of orogenic belts. Chemistry, Miyashiro, A., 1961. Evolution of metamorphic belts. Journal of Petrology 2, 277- 1-22. 311. Suzuki, K., Maruyama, S.,Yamamoto, S., Omori, S., 2010. Have the Japanese Islands Nakae, S., 1993. Jurassic accretionary complex of the Tamba Terrane, Southwest Japan, and its formative process. Journal of Geoscience, Osaka City University Journal of Geography 119, 1173-1196. 36, 15-70. Tagiri, M, Dunkley, D.J., Adachi, T., Hiroi, Y., Fanning, C.M., 2011. SHRIMP dating of Nakajima, T., 1994. The Ryoke plutonometamorphic belt: crustal section of the magmatism in the Hitachi metamorphic terrane, Abukuma Belt, Japan: Cretaceous Eurasian continental margin.Lithos 33,51-66. evidence for a Cambrian volcanic arc.The Island Arc 20, 259-279. Nakajima, T, 1996. Cretaceous granitoids in SW Japan and their bearing on the Tagiri, M., Morimoto, M., Mochizuki,R., Yokosuka, A, Dunkley, D.J., Adachi, T., 2010. crust-forming process in the eastern Eurasian margin. Transactions of the Royal Occurrence of metagranite of Cambrian SHRIMP zircon age in Hitachi Society of Edinburgh,Earth Sciences 95,249-263. metamorphic rocks and their geology. Journal of Geography 119, 245-256 (in Nakajima, T, Ishiwatari, A., Sano, S., Kunugiza, K, Okamura, M., Kano, T., Sohma, T. Japanese with English abstract). Hayasaka, Y., 1992. Geotraverse across the Southwest Japan arc: an overview of Taira,A., Katto,J, Tashiro, M., kamura, M., Kodama, K, 1988. The Shimanto Belt n Shikoku, Japan - evolution of Cretaceous to Miocene Accretionary Prism. Nakama, T, Hirata, T, Otoh, S., Maruyama, S., 2010. The oldest sedimentary age Modern Geology 12, 5-46. 472 Ma (Latest Early Ordovician) from Japan: U-Pb zircon age from the Takahashi, N., Kodaira, S., Klemperer, S.L, Tatsumi, Y, Kaneda, Y., Suyehiro, K., 1994. Hitoegane formation in the Hida Marginal Belt. Journal of Geography 119, 270- Crustal structure and evolution of the Mariana intra-oceanic island arc. Geology 278 (in Japanese with English abstract). 35,203-206. Nishimura, Y., 1990. Sangun metamorphic rocks': terrane problem. In: Ichikawa, K, Takahashi, N., Suyehiro, K, Shinohara, M., 1998. Implications from the seismic Mizutani, S., Hara, I., Hada, S., Yao, A. (Eds.), Pre-Cretaceous Terranes in Japan, crustal structure of the northern Izu-Bonin arc. The Island Arc 7, 383-394. vol. 224. IGCP Publication, Osaka, pp. 63-79. Tamura, Y., 2011. Formation of continental crust at the Izu-Honshu Collision Zone. Nishimura, Y., 1998.Geotectonic subdivision and areal extent of the Sangun belt, inner zone of Southwest Japan. Journal of Metamorphic Geology 16, 129-140. Tamura, Y., Ishizuka, O., Aoike,K., Kawate,S.,Kawabata, H., Chang, Q.,Saito, S., Nishimura, Y., Itaya, T., Isozaki, Y., Kameya, A, 1989. Depositional age and Tatsumi, Y., Arima, M., Takahashi, M, Kanamaru, T, Kodaira, S., Fiske, R.S., 2010. metamorphic history on 220 Ma high P/T type metamorphic rocks of the Missing Oligocene crust of the Izu-Bonin arc: consumed or rejuvenated during inner zone of the Southwest Japan. Memoirs of the Geological Society of Japan collision? Journal of Petrology 51, 823-846. 33,143-166. Tanaka,T.and Hoshino,M.,1987.Sm-Nd ages of Oki metamorphic rocks and its Nishimura, Y., Shibata, K., 1989. Ocurrence of metagabbro and its K-Ar age in geological significance. Abstract of 94th Annual Meeting of Geological Society of "Sangun metamorphic belt". Memoir of the Geological Society of Japan 33, 343- Japan, p. 492 (in Japanese). 357. Toriyama, R., 1954. Geology of Akiyoshi, Part Il, Stratigraphy of non-calcareous Otoh, S., Simizu, M., Ujike,O., Hasbaator, Sasaki, M.,Abe, M., Yamakita, S.1999. groups developed around the Akiyoshi Limestone Group. Kyushu University Geotectonic structure of East Asia before the opening of the Sea of Jap an -Part Faculty of Science Memoirs, Series D 5, 1-46. 1 Pre-Tertiary geotectonic framework. INAS Research Annual, Institute for Toriyama, R., 1958. Geology of Akiyoshi, Part Il, Fusulinids of Akiyoshi. Kyushu Northeast Asian Studies, Toyama University, vol. 25, pp. 115166. University Faculty of Science Memoirs,Series D 7,1-264. Otsuka, T.,1988.Paleozoic-Mesozoic sedimentary complex in the eastern Mino Tsujimori, T., 1999. Petrogenesis of the Fuko pass high-pressure metacumulate from Terrane, central Japan, and its Jurassic tectonism. Journal of Geoscience, Osaka the Oeyama peridotite body,southwestern Japan: evidence for Early City University 31, 63-112. Paleozoic subduction metamorphism. Memoir of Geological Society of Japan Sakashima, T,Terada, K,Takeshita, T., Hayasaka, T., ano, Y,Hidaka, H, 1999.Early 52,287-302. Ordovician granitoids from the Hitachi metamorphic rocks in the South Tsujimori, T., 2002. Prograde and retrograde P-T paths of the Late Paleozoic Kitakami belt and the Ryuhozan metamorphic rocks in the Higo belt. Japan, Glaucophane Eclogite from the Renge Metamorphic Belt, Hida Mountains, Abstract of the 106th Annual Meeting of the Geological Society of Japan, p. 63. Southwest Japan. International Geology Review 44, 797-818. Sakashima, T., Terada, K., Takeshita, T., Sano, Y., 2003. Large-scale displacement Tsujimori, T., Itaya, T., 1999. Blueschist-facies metamorphism during Paleozoic along the Median Tectonic Line, Japan: evidence from SHRIMP zircon U-Pb orogeny in southwestern Japan: Phengite K-Ar ages of blueschist-facies dating of granites and gneisses from the South Kitakami and pale-Ruyoke belts. tectonic blocks in a serpentinite melange beneath Early Paleozoic Oeyama Journal of Asian Earth Sciences 21, 1019-1039. ophiolite. The Island Arc 8, 190-205. Sano, H., Kanmera, K, 1988. Paleogeographic reconstruction of accreted oceanic Tsukada, K., 2003. Jurassic dextral and Cretaceous sinistral movements along the rocks, Akiyoshi, southwest Japan. Geology 16, 600-603. Hida marginal belt. Gondwana Research 6, 687-698. K.Wakita/Journal of Asian Earth Sciences 72 (2013） 75-87 87 Tsukada, K., Koike, T., 1997. Ordovician conodonts from the Hitoegane area, Wakita, K., 2000b. Cretaceous accretionary-collisional complexes in central Kamitakara Village, Gifu Prefecture. Journal of Geological Society of Japan 103, Indonesia. Journal of Asian Earth Sciences 18, 739-749. 171-174. Wakita, K., Igawa, T., Takarada,S., 2009. Seamless Geological Map of Japan at a Scale Tsukada, K, Takeuchi, T., Kojima, S., 2004. Redefinition of the Hida Gaien belt. of 1:200.000, DVD Version. Digital Geoscience Map G-16, Geological Survey of Journal of the Geological Society of Japan 110, 640-658 (in Japanese with Japan, AIST. English abstract). WakitaK,Mtcalf,I, 2005.cean plat stratiraphy in East and outheast Asia. Uchino, T, Kurihara, T., Kawamura, M., 2005. Early Carboniferous radiolarians Journal of Asian Earth Sciences 24, 670-702. discovered from the Hyachine Terrane, Northeast Japan: the oldest fossil age Yamakita, S., Otoh, S., 1998. Reconstruction of the geological continuity between for the clastic rocks of accretionary complex of Japan. Journal of the Primorye and Japan before opening of the Sea of Japan. INAS Research Annual Geological Society of Japan 111, 249-252 (in Japanese with English 24, 1-16. abstract). Yamazaki, T., Takahashi, M., Iryu, Y., Sato, T, Oda, M., Takayanagi, H., Chiyonobu, S., Uchino, T., Kawamura, M., Gouzu, C., Hyogo, H., 2008. Phengite Ar40/Ar39 age of Nishimura,A,Nakazawa,N., oka,T, 2010.Philippine Sea Plate motion since garnet bearing pelitic schist pebble obtained from conglomerate in the Nedamo the Eocene estimated from paleomagnetism of seafloor drill cores and gravity cores. Earth, Planets and Space 62, 495-502. 317 (in Japanese with English abstract). Yanai, S., Aoki, K., Akahori, Y., 2010. Opening of Japan Sea and major Tectonic lines Wakita, K, 1988. Origin of chaotically mixed rock bodies in the Early Jurassic to of Japan: MTL, TTL and Fossa Magna. Journal of Geography 119, 1079-1124 (in Early Cretaceous sedimentary complex of the Mino Terrane, central Japan. Japanese with English abstract). Bulletin of Geological Survey of Japan 39, 675-757. Zamoras,LR.,Matsuoka,A.,2000.Early Late Jurassic radiolarians from the clastic Wakita,K. 200oa.Melanges of the Mino terraneMmoir of Geological ociety of unit in usuanga Island,orth alawan,hilinesience Rert f Nigt Japan 55, 145-163. University, Series E (Geology) 15, 91-109.
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Wakita (2013) - geology and tectonics of Japanese islands a review.txt
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Prograde Metamorphism of the Horokanai Ophiolite
in the Kamuikotan Zone, Hokkaido, Japan
by HIDEO ISHIZUKA
Department of Geology, Kochi University, Kochi 780, Japan
(Received 27 April 1984; in revised form 15 October 1984)
ABSTRACT
The Horokanai ophiolite is a segment of metamorphosed oceanic crust and upper mantle,
tectonically emplaced into the Kamuikotan Zone of Hokkaido, Japan. Metamorphic grade, ranging
from the zeolite facies (Zone A), through the greenschist facies (Zone B) and the greenschist-amphibolite
transitional facies (Zone Q, to the amphibolite and granulite facies (Zone D), increases progressively
downwards with zone boundaries subparallel to the ophiolite pseudostratigraphy. The granulite facies
rocks include both metagabbros and their underlying ultramafic rocks. Coexisting minerals from
several tens of samples covering all the mineral zones were analysed by means of an electronprobe
microanalyser; the results are presented, along with brief consideration of their compositional variation
with metamorphic grade. The facies series of metamorphism of the Horokanai ophiolite corresponds to
the low-pressure type with a temperature range of 100-750 °C, which is broadly comparable to that
inferred for ocean-floor metamorphism. The major difference is the presence of the granulite facies
rocks in the Horokanai ophiolite and its absence in ocean-floor metamorphism.
INTRODUCTION
AJong the axial region of the island of Hokkaido, Japan, the Kamuikotan zone extends
from south to north for a length of 320 km and is 20 km wide or less (Fig. 1). Ishizuka et al.
(1983a) have characterized it as a zone where two distinct types of metamorphic rocks,
i.e. a high-pressure subduction complex and a low-pressure meta-ophiolitic rock, were
tectonically mixed up, accompanied by abundant ultramafic rocks. The low-pressure
meta-ophiolitic rock is generally fragmented into incoherent blocks; however, in the
Horokanai area (Fig. 1) a relatively intact suite is exposed, and hence it is named the
Horokanai ophiolite. The geology and geochemistry of the Horokanai ophiolite have been
studied in detail by Asahina & Komatsu (1979) and Ishizuka (1980a, 1981), and its tectonic
relationship to the high-pressure subduction complex has been discussed by Ishizuka et al.
(1983a). These previous studies indicate that the Horokanai ophiolite represents a
tectonically emplaced portion of metamorphosed oceanic crust and upper mantle (an
obducted ophiolite). In this paper, the detailed metamorphic petrology of the Horokanai
ophiolite is presented, along with discussion of its comparative significance to ocean-floor
metamorphism.
GEOLOGICAL SETTING
The Horokanai area is underlain by the Kamuikotan metamorphic rocks, the Horokanai
ophiolite, the Yezo Group, and Tertiary volcanic rocks (Fig. 1). The Kamuikotan
metamorphic rocks represent a high-pressure subduction complex, belonging to the
jadeite-glaucophane type of Miyashiro (1961), and are structurally bordered by the
overlying, thrust-faulted, low-pressure ophiolite. Ishizuka et al. (1983a) have demonstrated
that the ophiolite was tectonically emplaced (obducted) on to the Kamuikotan complex
[Jourml of Pctrolou, Vd 26, Pmrt 2, pp. 391-417. 1985]Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
392 H. ISHIZUKA
EXPLANATION
|VVVVV| Neogene Volcanic Rocks
Cretaceous Yezo Group
|| | || Kamuikotan Metamorphic
II I II Rocks
— HOROKANAI OPHIOLITE —
Ori!jmi|/Radiolartan Chert»
£?*,
M
'*rVBasaltic Rocks
pillow lava
tuff
hyaloclastite
massive lava flow
banded
amphibolite
Gabbroic Rocks
massive
amphibolite
-<CCumulate Rocks
Ultramafic Rocks
duniteo o>•o o
___ Dl
O •
a
.. „ A Thrust Fault
(barbs on upper plate)
Fault
Lithological Contact
FIG. 1. Geological sketch map of the Horokanai ophiolite (modified after Ishizuka, 1980a). The thin arrow in the
inset points to the locality of the Horokanai ophiolite.
along a thrust plane well after the cessation of both the low-pressure and high-pressure
metamorphic events. Recent radiometric and fossil data of the ophiolite are consistent with
their interpretation. The metamorphic hornblendes of the ophiolite give 40Ar-39Ar ages of
176-186 Ma (Takigami & Ishizuka, in prep.), and the chert member yields a radiolarian
assemblage of Tithonian age (Ishizuka et al., 19836); both being older than the K-Ar ages ofDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 393
the Karauikotan complex (72-145 Ma for muscovites: Imaizumi & Ueda, 1981). The Yezo
Group consists of unmetamorphosed Cretaceous sediments characterized by intercalated
accumulation of submarine elastics, and is regarded as forearc-basin deposits. The ophiolite
is in fault-contact with the Yezo Group, but in the southwestern part of the study area, Igi et
al. (1958) reported that the Yezo Group lay unconformably on chert which is now interpreted
as one of the chert members of the ophiolite. These structural, lithological and age
relationships are similar in some aspects to those in western California where the Coast
Range ophiolite (low-pressure meta-ophiolite), the Franciscan complex (high-pressure
subduction complex), and the Great Valley Sequence (forearc-basin sediments) are
technically juxtaposed (e.g. Page, 1981).
The sequence of the Horokanai ophiolite is so folded and faulted as to hinder a
quantitative estimate of its original thickness. Especially in the 'Mixed-Up Zone' of Fig. 1,
rocks of all ophiolite members are mixed up during the emplacement of the ophiolite
(Ishizuka, 1980a). The general sequence of the ophiolite is, however, confirmed by detailed
field mapping (Asahina & Komatsu, 1979; Ishizuka, 1980a), and it is schematically shown
in Fig. 1.
The top of the Horokanai ophiolite is Tithonian radiolaria-bearing chert. This unit is
underlain by close-packed basaltic pillow lavas and a complex of basaltic tuff-hyaloclastite-
massive lava flow. The hyaloclastite contains many pillow fragments embedded in a
cogenetic tuffaceous matrix. Deeper in the sequence, the nature of the precursors is
obliterated by metamorphic recrystallization, but we found that banded amphibolite is
underlain by massive amphibolite. This amphibolite sequence may represent a complex of
basaltic tuff-hyaloclastite followed by gabbro. In the above sequence, except for the
uppermost chert, numerous dikes of dolerite occur as feeder channels for the extrusive
basaltic rocks, but the typical sheeted dike complex described elsewhere (e.g. Coleman, 1977)
is absent. Near the boundary between the basaltic and gabbroic rocks, i.e. between the
banded and massive amphibolites, dikes of plagiogranite occur. These dikes of dolerite and
plagiogranite are also metamorphosed to the same grade as the adjacent host metabasites.
Further down the sequence occur ultramafic rocks consisting of dunite at the top followed by
foliated harzburgite and then massive harzburgite, with the intercalation of minor cumulates
composed of olivine-gabbro, pyroxenite, wehrlite and dunite. The ultramafic rocks are often
accompanied by layers or lenses of olivine-gabbro and microgabbro.
The bulk rock chemistry (major and minor elements such as Cr, Ni, Zr, and Y) of
representative rocks along with the chemistry of relict spinel and clinopyroxene in pillow
basalts (Ishizuka, 1981) indicates that the Horokanai ophiolite was essentially a tholeiitic
gabbro-basalt association with young transecting dikes of dolerite and plagiogranite. The
basal harzburgite represents the depleted mantle protolith which gave rise to formation of
the basaltic melts. Inasmuch as this ophiolite has a close geochemical affinity with normal
abyssal tholeiites (Ishizuka, 1981), an oceanic environment over an island arc or oceanic
island setting is envisaged for the site of its generation. A deep sea eruption far away from a
sialic arc or continental margin is also suggested by general lack of visible vesicles and
terrigenous debris in the extrusive members.
PETROGRAPHY AND MINERAL ZONES
Mafic rocks
The mafic rocks of the Horokanai ophiolite display significant changes in texture (Fig. 2)
and mineral paragenesis (Fig. 3) in response to changing metamorphic grade. The following
mineral zones are distinguished in the ascending order of metamorphic grade; ZoneDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
H. ISHIZUKA
FIG. 2. Photomicrographs of the Horokanai mafic rocks (all in plane polarized light). The scale bar is approximately
0-5 mm in length. Abbreviations are ol = olivinc, ex = dinopyroxene, pi = plagioclase, ac = actinolite,
ep = epidote, ch = chlorite, ab = albite, og = oligodase, hb = hornblende, ox = orthopyroxene and il = ilmenite.
A: Pillow basalt of Zone A showing well-preserved igneous texture. Note that olivine phenocryst is now replaced by
chlorite, and plagioclase lath by albite and/or zeolite, but groundmass dinopyroxene is unaltered. B: Hyaloclastite of
Zone B showing nematoblastic texture defined by acicular to prismatic actinolite. C: Hyalodastite of Zone C
showing coexistence of albite and oligoclase, and also that of actinolite and hornblende. Note Becke line at grain
boundaries between two plagioclases. D and E: Clinopyroxene-free and -bearing amphibolites of Zone D,
respectively, both showing granoblastic texture. F: Hornblende-granulite of Zone D showing granoblastic texture.
A: zeolites + chlorite + albite + pumpellyite, Zone B: albite + chlorite + actinolite + epidote,
Zone C: albite + oligoclase + actinolite + hornblende + chlorite + epidote, Zone D: horn-
blende + calcic plagioclase + dinopyroxene + orthopyroxene. These four mineral zones are
areally mapped (Fig. 4A); an example of the accuracy of locating zone boundaries may be
inspected in Fig. 4B. Most importantly, the grade of metamorphism increases down the
ophiolite stratigraphy, which is valid in all the mapped area except for the southern area
where several shear zones develop disturbing the ophiolite stratigraphy (Fig. 1) as well as the
distribution pattern of mineral zones (Fig. 4A).Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 395
Zone A
The upper sequence of the basaltic rocks (mainly pillow lava) belongs to this zone. The
pillow lava retains igneous structure and texture, but its igneous minerals are partially to
pervasively replaced by metamorphic minerals. Plagioclase phenocrysts and laths are
replaced by zeolites or albite that are associated with minor pumpellyite or calcite; ophitic to
subophitic clinopyroxene by chlorite; olivine phenocrysts by chlorite with minor calcite
and/or pumpellyite; Cr-spinel included in pseudomorphs after olivine or plagioclase
phenocrysts by Cr-rich chlorite. Interstitial glass is altered to chlorite with disseminated
fine-grained sphene and Fe-Ti oxide dust Fractures and veins are wholly filled by zeolites,
chlorite, albite, pumpellyite, calcite, and rarely quartz. Basaltic interpillow matrix, set into a
re-entrant space of close-packed pillow pile, is extensively recrystallized to the metamorphic
minerals given above, and this matrix sometimes contains chloritized or palagonitized glass
shards.
^LITHOLOGY
metamorphic \
minerals \
chabazlte
laumontlte
wairakite
stilblte
analcime
natrolite
thomsonlte
pumpellyite
prehnite
chlorite
albite
epidote
Ca-plagloclase
actinolite
hornblende
clinopyroxene
orthopyroxene
quartz
calcite
sphene
apatite
Fe-TI oxides
ZONEBASALTIC ROCKS \ GABBROIC ROCKS
Pillow : Tuff. Hyaloclastite ' Banded ( Massive
Lava : & Massive Lava Flow ' Amphlbolite \ Amphibolite
—i .....
[
4
£8*^ L8* iWS*
A"
BI-
L
c— — — ^—
D
CS*CHABAZITE SUBZOHE: LS* LAUMONTITE SUB2ONE: WS*.WAIRAKITE SUBZOHE
FIG. 3. Mineral paragenesis for the Horokanai mafic rocks.
Zeolites are stilbite, chabazite, laumontite, wairakite, natrolite, analcime, and thomsonite,
as determined optically and by X-ray diffraction. This zone can be further divided into three
subzones by the sequential appearance of chabazite, laumontite and wairakite (Fig. 3); each
subzone is mapped along a traverse route (Fig. 4C):
Chabazite subzone
(A-Cl) chlorite + chabazite + analcime + thomsonite
(A-C2) chlorite + chabazite + analcime + stilbite
Laumontite subzone
(A-Ll) chlorite + laumontite + analcime + thomsonite
(A-L2) chlorite + laumontite + thomsonite + albiteDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
H. ISHIZUKA
Opx-free
ZONE D
Opx-bearing
• ZONE A Opx I'ee
QZONE B ^ZONE D
OChabaziteSubzone
•'. LaumontiteSubzone©:WairakiteSubzone
Fia. 4. Mineral zone map of the Horokanai mafic rocks (A), distribution of rocks with critical mineral assemblages
of each mineral zone (B), and distribution of rocks containing critical mineral assemblages of each subzone of
Zone A (Q.
(A-L3) chlorite + laumontite + pumpellyite + albite
(A-L4) chlorite + laumontite + albite + quartz
(A-L5) chlorite + analcime + natrolite + thomsonite
(A-L6) chlorite + analcime + pumpellyiteDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 397
Wairakite subzone
(A-W1) chlorite + wairakite + analcime + thomsonite
(A-W2) chlorite + wairakite + thomsonite + albite
(A-W3) chlorite + wairakite + pumpellyite + albite
(A-W4) chlorite + wairakite + albite + quartz
Sphene and Fe-Ti oxide, and sometimes calcite, occur as the accessories. Some of the chlorite
minerals in the assemblages (A-Cl), (A-C2), and (A-Ll) are mixed-layer smectite/chlorite
clay, as determined by X-ray diffraction.
The mineral assemblages listed above are critical to the zeolite facies (e.g. Miyashiro &
Shido, 1970). The pumpellyite-bearing assemblages also occur in other zeolite facies
metamorphic terrains (Boles & Coombs, 1977; Liou, 1979; Evarts & Schiffman, 1983).
Since chabazite, laumontite, and wairakite have the same composition except for H2O
(CaAl2Si4O12.nH2O: n = 6 for chabazite, n = 4 for laumontite, n = 2 for wairakite), the
change from the chabazite through the laumontite to the wairakite subzone represents
a progressive dehydration sequence of Ca-Al zeolites with rising temperatures. The
equilibrium dehydration of laumontite to wairakite+ 2H2O has been ascertained experi-
mentally by Liou (1971a). On the other hand, quartz occurring in the laumontite and
wairakite subzones coexists with albite but not analcime. This suggests that these subzones
are placed within the albite + quartz field, the higher-temperature field than the dehydration
of analcime + quartz (Liou, 1971b; Thompson, 1971).
In the highest-grade part of Zone A occur the zeolite-free assemblages prehnite +
pumpellyite + chlorite and prehnite + actinolite + chlorite. The former assemblage is critical
to the prehnite-pumpellyite facies, and the latter one has been reported from the Karmutsen
contact metamorphic aureole of Vancouver Island (Kuniyoshi & Liou, 1976) and from the
East Taiwan ophiolite (Liou & Ernst, 1979) and may belong to the prehnite-actinolite facies
of Liou et al. (in press). Both assemblages occur only in the well-recrystallized interpillow
matrix; the nearby pillow core retains relict clinopyroxene and contains the chlorite+
wairakite + pumpellyite assemblage which is critical to the wairakite subzone.
Zone B
This zone includes the middle sequence of the basaltic rocks (mainly hyaloclastite) and is
characterized by the actinolite + chlorite+ epidote +albite assemblage. There is a difference
in texture between the hyaloclastite matrix and enclosed pillow fragments. The matrix is well
recrystallized to yield a nematoblastic texture defined by actinolite. The pillow fragments
commonly retain igneous texture, but plagioclase is totally albitized with trace epidote, and
clinopyroxene is extensively uralitized to aggregates of chlorite and actinolite. Furthermore,
the modal proportion of metamorphic minerals depends upon the nature of the original
rocks; actinolite tends to occur more abundantly in the matrix than in the pillow fragments.
Both the matrix and pillow fragments contain various amounts of quartz, sphene, Fe-Ti
oxide, and rarely calcite.
The ubiquitous distribution of actinolite, chlorite, epidote, and albite indicates that
Zone B belongs to the greenschist facies. In the lowest-grade part of this zone, the epidote-
free assemblage pumpellyite + actinolite + chlorite + albite occurs, but it is too restricted to
define an areal mineral zone.
Zone C
This zone occurs occupying the middle to lower sequence of the basaltic rocks (mainly the
lowermost part of hyaloclastite, and the uppermost part of banded amphibolite). TheDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
398 H. ISHIZUKA
hyaloclastite exhibits nematoblastic texture defined by actinolite and/or hornblende,
whereas the shape of pillow fragments is commonly transitional into the nematoblastic
matrix. The banded amphibolite shows two distinct bands of less than 05 mm width; one rich
in fine-grained plagioclase (albite and/or oligoclase) and the other in nematoblastic
hornblende. Other constituent minerals include chlorite, epidote, sphene, Fe-Ti oxide, and
rarely quartz and calcite.
In most of the Zone C rocks, albite and oligoclase coexist as discrete grains of less
than 0-3 mm diameter, which often gather to form pool-shaped aggregates of less than
20 mm diameter. The grain boundary between albite and oligoclase is commonly sharp,
and straight or gently curved; a Becke line is sometimes distinct at the grain boundary.
Occasionally, oligoclase rims of less than 50 /zm width develop around a few of the albite
grains. Also, in the medium-grade part of Zone C, actinolite and hornblende coexist as
discrete acicular to prismatic crystals, being less than 0-5 mm in length, which define a
lineation. Uncommonly, actinolite lamellae of about 20 urn in width occur in a host
hornblende. The textural relationships characterized by discrete grains or crystals may
indicate equilibrium coexistence of albite with oligoclase, and actinolite with hornblende,
respectively. It is, however, uncertain whether the core-rim relationship of albite and
oligoclase and the lamella intergrowth of actinolite and hornblende represent equilibrium or
disequilibrium textures.
The optical differentiation of two amphiboles and sometimes two plagioclases is very
difficult, and the mineral assemblages in this zone were determined by using the electron
microprobe; they are listed below:
(C-1) actinolite + albite + oligoclase + chlorite + epidote + quartz
(C-2) actinolite + albite -I- oligoclase + chlorite
(C-3) actinolite + oligoclase+chlorite
(C-4) actinolite + hornblende + albite + chlorite+epidote
(C-5) actinolite + hornblende + albite + oligoclase + chlorite -I- epidote + quartz
(C-6) actinolite + hornblende + albite + chlorite
(C-7) actinolite + hornblende + albite + oligoclase + chlorite + quartz
(C-8) hornblende + albite + oligoclase + chlorite
(C-9) hornblende + albite + oligoclase
Sphene and Fe-Ti oxide are the common accessory minerals, but calcite is confined to the
assemblages (C-4) and (C-6). One of the most characteristic features of these assemblages is
the coexistence of albite and oligoclase, and hence we may call this zone the peristerite zone.
With respect to the metamorphic fades, the assemblage actinolite + oligoclase + chlorite
(C-3) is critical to the actinolite-caldc plagioclase fades of Miyashiro (1961), which has been
reported from the transitional zone between the greenschist and amphibolite fades in contact
metamorphic aureoles (e.g. Abukuma Plateau: Shido, 1958; Kitakami Mountains: Seki, 1961;
Sierra Nevada: Loomis, 1966; Yap Islands: Shiraki, 1971; Vancouver Island: Kuniyoshi &
Liou, 1976) as well as from the Mid-Atlantic Ridge (Miyashiro et al., 1971) and from the East
Taiwan ophiolite (Liou & Ernst, 1979). However, the majority of the Zone C rocks contain
the peristerite pair with or without the actinolite-hornblende pair, and in a strict sense they
do not belong to the actinolite-calcic plagioclase fades. In this connection, it should be noted
that the use of the electronprobe microanalyser has recently revealed the common
occurrence of the peristerite and/or actinolite-hornblende pairs in the transitional zone of
contact metamorphic aureoles (e.g. Abukuma Plateau: Tagiri, 1973, 1977; Sierra Nevada:
Hietanen, 1974; Yap Islands: Maruyama et al., 1982, 1983). Such finding requires a
re-examination of the mineral assemblages to define the actinolite-caldc plagioclase fades.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 399
Therefore, in this paper, the Horokanai Zone C is referred to simply as the transitional facies
from the greenschist to amphibolite facies.
Zone D
Zone D includes the lowermost sequence of the basaltic rocks (banded amphibolite), and
all the gabbroic rocks (massive amphibolite). Both banded and massive amphibolites display
granoblastic to gneissose texture, but trace igneous clinopyroxene is preserved in the massive
amphibolite (metagrabbro) as porphyroclastic grains with margins and/or local patches
recrystallized to fine-grained clinopyroxene and/or hornblende. The alternation of leuco-
cratic (plagioclase-rich) and melanocratic (hornblende-rich) bands, measuring less than 10
cm thick, is typical for a banded amphibolite.
The constituent minerals include calcic plagioclase, hornblende, clinopyroxene, ortho-
pyroxene, sphene, Fe-Ti oxide, and rarely apatite, to which chlorite is added in the
lowermost-grade part of this zone. Calcic plagioclase is usually saussuritized to aggregates of
fine-grained albite ± chlorite ± epidote + prehnite. The Z-axial colour of hornblende changes
from blue-green to brown with increasing metamorphic grade. Metamorphic clinopyroxene
first appears in the medium-grade part of Zone D, and its modal proportion tends to increase
gradually with increasing metamorphic grade. In the highest-grade part of this zone,
metamorphic orthopyroxene develops with conspicuous pleochroism varying from salmon
pink to pale green.
The observed mineral assemblages in this zone are:
(D-l) calcic plagioclase + hornblende + chlorite
(D-2) calcic plagioclase + hornblende
(D-3) calcic plagioclase + hornblende+clinopyroxene
(D-4) calcic plagioclase + hornblende + clinopyroxene-I-orthopyroxene
Fe-Ti oxide is common, but sphene is confined to the lower- to medium-grade part of this
zone. Uncommonly, apatite occurs in the assemblage (D-3).
The widespread occurrence of the calcic plagioclase + hornblende assemblage indicates
that most of the Zone D rocks belong to the amphibolite facies, but the highest-grade rocks
with pleochroic orthopyroxene may be placed in the (hornblende-) granulite facies (e.g.
Howie, 1965).
Ultramafic rocks
The ultramafic rocks of the Horokanai ophiolite are serpentinized to some degree, but
their parental peridotite types are inferred from relict mineralogy to have been harzburgite
and dunite with minor pyroxenite.
Less serpentinized ultramafic rocks have curvilinear grain boundaries between olivine
crystals or between olivine and orthopyroxene crystals; the grain size of the olivine and
orthopyroxene is typically coarse. Spinel and clinopyroxene, of which the latter is confined to
the upper horizon of the ultramafic rocks, are rather smaller in grain size. This kind of
peridotite texture is characteristic of the protogranular type, one of the representative
textures of the mantle peridotite (Merrier & Nicolas, 1975). However, there is little sign of
polygonization or recrystallization of large crystals into aggregates of finer grains with
preferred orientation, even though coarse-grained olivine and orthopyroxene sometimes
exhibit wavy extinction and kink-bands. Such a textural relationship indicates that the
Horokanai ultramafic rocks have not undergone intense deformation or plastic flow in the
mantle (Merrier & Nicolas, 1975).Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
400 H. ISHIZUKA
The main features of serpentinization include: (1) olivine is split into fine, rounded
grains floating in a network of serpentine and magnetite, (2) orthopyroxene is almost
entirely altered to bastite, (3) spinel with translucent brown colour is little altered, but
its margin is sometimes armoured by Cr-rich chlorite, and (4) clinopyroxene is relatively
fresh with a trace of chlorite rinds. In addition to these altered minerals, talc and
calcite occur, along with soda-tremolite as described by Ishizuka (1980b). Serpentine
minerals are chrysotile and lizardite, and no antigorite is detected in the analysed
samples except one sample that contains antigorite cross-cut by a veinlet of chrysotile
or lizardite.
Olivine, pyroxenes and spinel described above are primary minerals once equilibrated
with some high-temperature magmas, but their chemical compositions, as will be described
in a later section, have been substantially modified from the high-temperature stage through
element redistribution (re-equilibration) at temperatures corresponding to the granulite
facies.
TABLE 1
Mineral assemblages of analysed samples from the Horokanai ophiolite
Sample
no.
AC151
AL351
AL331
AL651
AW151
AW351
AW331
B1501
B15O2
B15O3
C1501
C2501
C4301
C5501
C6501
C73O1
C8501
C9501
C93O1
D1501
D2501
D2502
D2503
D3501
D35O2*'
D33O1
D4301
D4302
D4401Zone
A
A
A
A
A
A
A
B
B
B
C
C
C
C
C
C
C
C
C
D
D
D
D
D
D
D
D
D
DZe
X
X
X
X
X
X
X
—
—
—
—
—
—
—
—
—
—
—
—
_
—
—
—
—
—
_
—
—Pu
X
X
X
—
X
X
—
—
—
—
—
—
—
—
_
—
—
—
—
—
—
—
—
—
—
—
—Ch
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
—
—
X
—
—
—
—
—
—
_
—Ab
X
X
—
—
X
X
X
X
X
X
X
X
X
X
X
X
X
X
—
—
—
—
—
—
—
—
—Ep
—
—
—
—
—
X
X
X
X
—
X
X
—
—
—
—
—
—
—
—
—
—
—
—
_
—
—PI
—
—
—
—
—
—
—
—
X
X
—
X
—
X
X
X
X
X
X
X
X
X
X
X
X
X
XAt
—
—
—
—
X
X
X
X
X
X
X
X
X
—
—
—
—
—
—
—
—
—
—Hb
—
—
—
—
—
—
—
_
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
XCx
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
X
X
X
X
X
XOx
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
—
X
X
X—
—
—
—
X
X
X
X
—
—
X
X
—
—
—
—
—
—
—
—
—
—Cc
X
—
_
X
—
_
—
X
—
—
—
X
—
X
—
—
—
—
—
—
—
—
—
—
—Sp
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
—
—
—
—
—FT
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
XF*/M
112
1-21
1-30
_
—
1 29
1-02
—
098
—
117
1-22
1-22
1-30
—
—
—
099
—
—
O96
—
085
—
065
063
Abbreviations: Ze = zeolites, Pu = pumpellyite, Ch = chlorite, Ab = albite, Ep = epidote, PI = Ca-plagioclase,
At = actinolite, Hb = hornblende, Cx = clinopyroxene, Ox = orthopyroxene, Qz = quartz, Cc = calcite, Sp •=
sphene,FT = Fe-Ti oxide, F*/M = FeO*/MgO bulk rock ratio quoted from Ishizuka(1981)(FeO* - total iron as
FeO), x = present, — = absent. D35O2** contains small amounts of apatite.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE
MINERAL CHEMISTRY401
Chemical compositions of selected minerals were determined using the Hitachi Model
XMA-5A electronprobe microanalyser, at Kanazawa University. The Bence-Albee (1968)
method with correction factors of Albee & Ray (1970) was used for data reduction.
Supplementary microprobe analyses were done using the JEOL Model JAX-50A, at
Hokkaido University, and the JEOL Model J AX-5A, at Kochi University; both microprobes
using the same data correction procedure as that of the Hitachi microprobe. Mineral
assemblages of analysed samples are given in Table 1, along with the FeO*/MgO bulk rock
ratio quoted from Ishizuka (1981). Representative analyses of selected minerals are listed in
Tables 2 to 9.
TABLE 2
Representative analyses of zeolite species from Horokanai Zone A
Sample no.
SiO2
A12O3
Fe2O?
CaO
Na2O
K.2O
Total
Si
Al
Fe
Ca
Na
KChabazite
AC151
4713
18-89
008
1085
041
005
77-41
4O45
1-911
0O05
0998
O068
0005Laumontite
AU51
5081
21-79
010
11-85
032
003
84-90Wairakite
AW351
54-85
23O2
013
1212
002
002
9O16
Atomic proportions*
3-977
2011
O006
O994
0049
O0034024
1-991
0007
0953
0003
OO02Analcime
AL651
54-21
22-87
007
009
13-75
0O3
9102
2O03
0996
0002
0004
0985
0O01Ca-analcime
AW 151
54-87
22-96
005
3-96
9-82
0O2
91-68
2O05
O989
O001
0155
0696
O001Thomsonite
AC151
37-21
3210
005
1402
3-97
002
87-37
4-957
5O42
O005
2O01
1026
0O03
Fe2O* means total iron as Fe2O3.
Atomic proportions" were calculated on the basis of O = 12 for chabazite, laumontite, and
wairakite, O = 6 for analcime, and O •= 20 for thomsonite.
Zeolites (Table 2)
Zeolites occurring in Zone A are commonly homogeneous in composition, approaching
the ideal stoichiometry of chabazite, laumontite, wairakite, analcime, and thomsonite,
respectively. However, analcime occurring in the wairakite subzone as a pseudomorph after
plagioclase has a CaO content of up to 4 wt. per cent, and shows substantial solid solution
toward wairakite. Such a calcium-bearing analcime is rare, but has been described in
low-grade metamorphic terrains (Seki, 1971; Surdam, 1973; Evarts & Schiffman, 1983).
Pumpellyite (Table 3)
Pumpellyite occurs in the laumontite and wairakite subzones of Zone A; the com-
positional variation in terms of Al-Fe*-Mg is illustrated in Fig. 5A.
Using the formula for pumpellyite proposed by Coombs et al. (1976): W4X2Y4Z6O(20+i)
.^ where W = (Ca, Mn); X = (Mg, Fe2 + , Mn^.^Fe3*, Al),,; Y = (Al, Fe3+);Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
402 H. ISHIZUKA
TABLE 3
Representative analyses of pumpellyites from Horokanai Zone A
Sample no.
SiOj
A12O3
FeO»
MnO
MgO
CaO
Total
Si
Al
Fe
Mn
Mg
CaAL351
36-31
13-57
19-89
0-16
2-61
22-31
94-85
Atomic
5-998
2-642
2-747
0O23
0-642
3-948AI331
36-28
14-64
18-71
0-13
2-24
22-39
94-39
proportions,
6-008
2-858
2-591
0-018
0-553
3-973AL651
36-04
15-76
16-33
0-15
2-59
22-34
93-21
total cations
5-994
3-090
2-271
O021
0-642
3-981AW351
36-34
17-55
14-36
0-13
2-43
22-49
93-30
= 16-00
6003
3-417
1-984
0-018
0-598
3-980AW331
35-87
18-92
13-36
0O9
2-39
21-67
92-30
5-965
3-710
1-858
0-013
0-593
3-862
AL351, AL331, and AL651 are from the laumontite subzonc, and AW351
and AW331 from the wairakite subzone.
FeO* means total iron as FeO.
AlPUMPELLYITE
Fe1• : Laumontlte
Subzone
o - Wairakite
Subzone
60 50CHLORITE
20
50 SO
10
>t
oc
5O"
0
L
n+
ZONE C+
ZONE B+ +
EPIDOTE
0.1 0.2 0.3 0.4
FIG. 5. Compositional variations of Horokanai pumpellyites (A), chlorites (B), and epidotes (Q.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOL1TE 403
Z = (Si, Al), we can depict several features. The Si value in the Z-site (5-97-6-01) is close to the
ideal 6. The Fe3+ value in the Y-site (0-33-1-36) is high, and so Al is absent in the X-site. It is
also clear that Fe* (total iron) varies antithetically relative to Al (Fig. 5A), indicating that the
observed variation is mainly due to Fe3 + ^± Al substitution in the Y-site. The X-site
(2-00-2-13) is in excess of the ideal value of 2, while the available Ca and Mn in the W-site
(3-88 — 400) cannot fill the ideal figure of 4. The total occupants of the X- and W-sites nearly
equal 6, suggesting that other divalent cations, namely Fe2 + or Mg in the X-site, may also
enter the W-site. Such pumpellyite chemistry as shown above may be most characteristic
of zeolite fades metamorphic terrains (Boles & Coombs, 1977; Liou, 1979; Evarts &
Schiffman, 1983).
On the other hand, it is noteworthy that the pumpellyites from the wairakite subzone are
TABLE 4
Representative analyses of chlorites from Horokanai Zones A, B, C, and D
Sample no.
SiOj
A12O3
FeO*
MnO
MgO
CaO
Total
Si
Al
Fe
Mn
Mg
Ca
Sample no.
SiOj
AljOj
FeO1
MnO
MgO
CaO
Total
Si
Al
Fe
Mn
Mg
CaAC151
29-21
16-55
23-77
0-30
1815
008
88-06
6-061
4-049
4124
0052
5-612
0017
C1501
27-89
18-56
19-67
0-32
21-51
008
88-03
5-681
4-457
3-351
0055
6-530
0017AL351
29-56
1603
24-35
029
17-91
009
88-23
6137
3-923
4-228
0051
5-542
0020
C2501
28-29
1805
2O85
017
21-29
O10
88-75
5-743
4-320
3-540
O029
6-441
0022Zone A
AL331
29-30
16-51
24-25
023
17-95
012
88-36
Atomic
6O70
4033
4-202
0040
5-542
0026
C4301
28-45
18-32
18-74
038
22-22
012
88-23
Atomic
5-750
4-365
3168
0066
6-693
0026AL651
3O14
15-57
22-41
031
19-70
008
88-21
proportions, 0
6193
3-772
3-851
0054
6033
0017
Zone C
C5501
28-81
17-61
18-79
030
22-82
008
88-41
proportions, 0
5-811
4187
3169
0051
6-860
0017AW351
29O6
1711
2305
028
18-67
007
88-24
= 28-0
5-993
4159
3-975
0048
5-737
0015
0550/
29-43
17-21
1819
033
23-47
Oil
88-74
•=280
5-890
4O61
3045
0057
7O01
0024B1501
28-65
16-82
22-75
026
19-51
004
8803
5-924
4100
3-934
0046
6O13
0009
C7301
29-40
17-85
16-89
031
23-70
009
88-24
5-873
4-204
2-822
0053
7056
O019Zone B
BI502
28-39
1710
21-62
027
2O69
0O7
8814
5-835
4143
3-716
0O47
6-337
0015
C8501
29-29
17-93
16-11
027
24-52
007
8819
5-833
4-210
2-683
0045
7-277
0014BIS03
28-65
17-20
21-22
035
2059
O10
8811
5-875
4158
3-638
0060
6-292
0022
Zone D
D1501
29-15
17-99
15-89
025
24-73
008
8809
5-807
4-225
2-648
0042
7-342
0017
FeO* means total iron as FeO.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
404 H. ISHIZUKA
distinctly higher in Al than those from the laumontite subzone (Fig. 5A). This suggests that
the Al content of pumpellyite may be temperature dependent; the details will be given
elsewhere (Ishizuka, in prep.).
Chlorite (Table 4)
Chlorite is abundant in Zones A, B, and C, but rare in Zone D; the compositional variation
in terms of Al-Fe*-Mg against metamorphic grade is illustrated in Fig. 5B.
Chlorite is generally restricted in composition within individual samples, except for
mixed-layer smectite/chlorite clay in Zone A that has the compositions varying considerably
even within a single sample, e.g. SiO2 = 35-42 wt. per cent, A12O3 = 12-17 wt. per cent,
FeO* = 15-27 wt. per cent, MgO = 10-14 wt. per cent. The Fe*/(Fe* + Mg) ratio in chlorite
decreases systematically with increasing metamorphic grade, ranging from 043 (Zone A) to
027 (Zone D). Since the FeO*/MgO bulk rock ratio in chlorite-bearing rocks ranges from
098 to 1 -30 but shows no sign of systematic variation against metamorphic grade (Table 1), it
is suggested that the Fe*/(Fe* + Mg) ratio in chlorite tends to decrease with increasing
metamorphic grade. Similar compositional trends have been described in other meta-
morphic terrains (e.g. Cooper, 1972; Kurata & Banno, 1974). On the contrary, the
SiO2/(SiO2 + A12O3) ratio in chlorite is nearly constant throughout all the mineral zones
(0-63 ±0-03), which contradicts a tendency that this ratio decreases with increasing
metamorphic grade, as described in many metamorphic terrains (cf. fig. 3 of Maruyama
et al, 1983).
Epidote (Table 5)
Epidote is common in Zone B, but sporadic in Zone C; the frequency distribution of
Xp, = Fe3+/(Fe3+ + Al) is illustrated in Fig. 5C.
In Zone B, epidote is usually zoned with Xp, decreasing from the core to the rim.
TABLE 5
Representative analyses of epidotes from Horokanai Zones B and C
Sample no.
SiOj
TiO2
A12O,
Fc2Of
MnO
MgO
CaO
Total
Si
Ti
Al
Fe
Mn
Mg
CaB1501
core-rim
37-58
001
22-95
13-77
015
0O4
23-33
97-83
3-002
0001
2161
0828
0010
O005
1-99737-95
003
25-46
1052
017
002
23-59
97-74
2-997
0002
2-370
0625
0011
0002
1-996Zone B
B1502
core-rim
37-39
005
23-42
12-97
023
O07
23-16
97-293810
O02
26-18
9-41
015
002
23-68
97-56BI503
core-rim
37-74
002
24O2
12-41
O20
0O9
23-41
97-89
Atomic proportions, 0 =
2-996
0003
2-212
O782
O015
O008
1-9883001
O001
2-431
0558
O010
0002
1-9992-998
0O01
2-249
0742
0013
0011
1-9923816
003
26-98
8-27
021
012
23-65
97-42
12-5
2-998
OO02
2-499
0489
0014
0014
1-991C1501
38-48
002
28-35
6-70
O10
003
23-75
97-43
3-001
O001
2-607
0393
O007
0O03
1-985Zone C
C4301
38-39
O03
28-92
6O5
012
002
23-85
97-38
2-990
OO02
2-655
0355
OO08
0002
1-990C550I
38-45
002
28-99
6O0
Oil
003
23-88
97-48
2-991
0001
2-659
O351
0007
0003
1-990
Fe2OJ means total iron as Fe2O3.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 405
Comparing samples from the low-grade part (no. Bl 501) and the high-grade part (no. B15O3)
in Zone B, the minimum Xp, (i.e. Xp, of the rim) is higher in the former (Xp, = 021) than the
latter (Xp, = 0-16). Both samples have the same mineral assemblage (epidote + actinolite +
chlorite + albite + quartz + sphene) and a similar Fe*/(Fe* + Mg) ratio in chlorite (0-38 ±
0-02). The identified opaque phases include ilmenite, magnetite and pyrite, but no hematite.
Such a relationship between Xp, and metamorphic grade is consistent with the observation of
Miyashiro & Seki (1958), and with the model system of Nakajima et al. (1977), that is, the
minimum Xp, of the epidote + chlorite + actinolite assemblage is temperature dependent and
decreases with increasing metamorphic grade. On this basis, chemical zoning of Zone B
epidotes is interpreted to represent the prograde stage of metamorphism.
In Zone C, epidote is relatively homogeneous in composition, and has lower Xp, (0-08-
0-15) than that of Zone B.
Plagioclase (Table 6)
Plagioclase is ubiquitous in all the mineral zones; the variation of the An-content against
metamorphic grade is illustrated in Fig. 6A.
TABLE 6
Compositions of plagioclases from Horokanai Zones A, B, C, and D
Sample
no.
AL351
AL331
AW351
B1501
B15O2
B15O3
C1501**
C2501**
C4301
C55O1"
C65O1
C7301**
C8501'*
C9501**
C9301**
D1501
D25O1
D25O2
D25O3
D35O1
D35O2
D33O1
D4301
D43O2
D4401Zone
A
A
A
B
B
B
C
C
C
C
C
C
C
C
C
D
D
D
D
D
D
D
D
D
DAn-content
range
05-2-0
07-2-2
08-3-0
08-4-5
1-0-4-7
1-4-6-3
{ 1-8-7-0
1 201-24-7
{ 1-3-8-5
1 18-9-22-8
1-9-9-5
{ 1-8-7-5
1 17-6-22-3
1-5-5-8
{ 1-3-5-0
1 15-5-19-8
{ 1-1-3-8
1 13-4-16-6
{ 07-3-7
1 101-14-5
{ 05-2-1
1 8-0-11-3
9-2-203
18-9-35-4
27-6-41-4
36-6-4O2
44-3-5O2
46-9-531
602-64-5
62-4-69-3
75-0-82-5
77-6-89-2average
09
1-0
1-3
20
2-3
4-2
f 4-6
122-3
{ 6O
121-2
6-4
I 19-3
40
f 31
1 173
f 2-5
1 15O
1-9
112-2
( 1-2
1 9-7
15-7
23-9
33-8
38-7
48-8
5O6
62-9
65-3
77-2
85-4K2O
wt. per cent
average
009
012
013
008
020
010
{ 007
I 0-12
{Oil
I 0-13
005
{006
1009
017
{ 013
1024
{012
1 013
{021
IO20
{010
1 014
027
012
019
007
015
029
022
017
012
020Fe2O$
wt. per cent.
average
009
027
022
025
037
010
{021
1024
{037
1 039
024
{017
I Oil
007
{043
1045
{031
1O20
{Oil
I 012
{008
1 009
029
029
035
023
O09
017
O33
043
039
044Number of
analyses
9
9
12
15
15
15
{49
I 9
{ 23
1 6
17
{21
I 9
16
{ 17
I 10
f 10
\ 5
{ 13
1 713
1 9
24
18
12
9
13
15
11
19
13
19
Fe2OJ means total iron as Fe2O3. Samples with asterisks (**) contain two-phase plagioclases.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
406 H. ISHIZUKA
Sample
No.
04401
D4302
D43O1
03301
03502
D35O1
D25O3
02902
D2S01
D1S0
~C»301
C8501
CSS01
C73O1
C6S01
CSS01
C43O1
C25O1
.C150L
81503^
B1S02
_B1501_
AW3 51
AL331
AL3B110An-Content
30 50PLAGIOCLASE
70 00ZONE
B
A
Sample
No.
04401
04302
04301
D3301
O3SO2
D3S01
D2SO3
D2SO2
D2B01
D1S01
"C9301"
CB501
C8501
C73O1
C6S01
C9S01
C4301
C2901
C1S01
"B1803"
B1S02
B1S01AMPHIBOLE
0.5 1.5ZONE
FIG. 6. Variations of An-content (A) and A\n (B) for Horokanai plagioclases and amphibolcs, respectively. Sample
No. is arranged from the bottom to the top in the ascending order of metamorphic grade.
In Zones A and B, plagioclase is albite, and its average An-content increases with
increasing metamorphic grade, ranging from 0-9 per cent in the lower-grade part of Zone A
(no. AL351) to 4-2 per cent in the higher-grade part of Zone B (no. B15O3).
In Zone C, both albite and oligoclase (peristerite pair) coexist as discrete grains (Fig. 2Q,
and a distinct compositional gap can be denned. With increasing metamorphic grade, the
average An-content in albite first increases from 4-6 (no. C1501) to 60 (no. C2501) and then
decreases from 5-0 (no. C5501) to 1-2 (no. C9301), whereas the average An-content in
coexisting oligoclase decreases successively from 22-3 (no. C1501) to 9-7 (no. C9301). As
shown in Table 6 by the relative amounts of coexisting albite and oligoclase in individual
samples, increasing metamorphic grade appears to result in the formation of greater amounts
of oligoclase in comparison to albite. A preliminary study on a few grains shows that the
albite grains are relatively uniform in composition, but two types of compositional zoning are
detected in the coexisting oligoclase grains: one has a sodic margin, and the other a calcic one,
the width of the margin being less than 10 //m in both types. Orville (1974) has considered thatDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 407
the peristeritc gap is a two-phase binary loop involving crystallographic transformation in
albite-rich compositions. However, many descriptions of natural rocks have shown that the
peristerite gap is an asymmetrical solvus delineated by a steep sodic limb and a gentle calcic
limb (e.g. fig. 7 of Maruyama et al., 1982) or by a gentle sodic limb and a steep calcic limb
(Tagiri, 1973). As is clearly shown in Fig. 6A, there is a compositional shift of the Horokanai
peristerite gap toward the lower An-content with increasing metamorphic grade. This does
not support the concept of a solvus and is favourable to a transitional loop, but certainly
more data are needed to substantiate this relationship. On the other hand, the calcite-bearing
samples from Zone C (nos. C4301 and C6501) contain albite but are devoid of oligoclase,
suggesting that PQQ^ may be a factor that plays a role in inhibiting oligoclase-producing
reactions (Cooper, 1972).
In Zone D, one-phase plagioclase occurs. Although a large range of the An-content is
encountered for each sample, the average An-content again increases with increasing
metamorphic grade, ranging up to bytownite.
All the analysed plagioclases are extremely low in Fe2Of and K2O, less than 05 and 03
wt. per cent, respectively.
Amphibole (Table 7)
Amphibole occurs in Zones B, C, and D; the variation of A1IV against metamorphic grade
is illustrated in Fig. 6B, in which the double circles pointed to by Table T nearly correspond
to the maximum A11V frequency within individual samples, and the plot of (Na + K) and
Fe*/(Fe* + Mg) versus A1IV is shown in Fig. 7.
In Zone B to the lower-grade part of Zone C, amphibole is actinolite, and its maximum
A1IV frequency increases gradually with increasing grade of metamorphism, ranging from
010 (no. B1501) to 0-50 (no. C2501).
In the medium-grade part of Zone C, both actinolite and hornblende coexist as discrete
acicular to prismatic crystals (Fig. 2C), and a distinct compositional gap can be defined. With
increasing metamorphic grade, the maximum A1IV frequency in actinolite increases from 0-61
(no. C4301) to 0-82 (no. C6501), whereas that in coexisting hornblende decreases from 1-27
(no. C4301) to 104 (no. C6501), and finally both converge at the crest about A1IV = 0-9. It
seems that, with increasing metamorphic grade, the increase in A1IV of actinolite is not as
rapid as the decrease in A1IV of hornblende, and a slightly asymmetrical gap results. The
Fe*/(Fe* + Mg) ratio is commonly lower in actinolite than its coexisting hornblende.
Inasmuch as the FeO*/MgO bulk rock ratio of two-amphibole-bearing rocks is nearly
constant (1-22-1-30: Table 1), its effect on the compositional gap is presumably small. Tagiri
(1977) and Maruyama et al. (1983) described the compositional gap between actinolite and
hornblende as shifting toward the lower A1IV content with increasing metamorphic grade,
and suggested that such a compositional shift was attributed to a transitional loop associated
with a crystallographic transformation in actinolite-rich composition. The compositional
gap as shown in Fig. 6B shows no sign of such a compositional shift. In sample C7301 which
occurs in the higher-grade part of Zone C than sample C6501, actinolite coexisting with
hornblende has a maximum A1IV frequency of 0-79 lower than that of sample C6501, and then
the actinolitic limb of the gap appears to bend toward the lower A1IV content. However, this
actinolite is not a discrete crystal but a lamella in a host hornblende. Comparing only discrete
crystals of coexisting actinolite and hornblende, the compositional gap delineated in the
present study, especially the successive change of actinolitic limb compositions, favours a
solvus over a transitional loop, but this needs to be corroborated still more.
In the higher-grade part of Zone C to the lower-grade part of Zone D, amphibole is
hornblende, but its A1IV content varies considerably from one rock to another. Such variationDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
TABLE 7
Representative analyses of amphiboles from Horokanai Zones B, C, and D
Sample no.
SiOj
TiOj
AI2O3
FeO*
MnO
MgO
CaO
Na2O
K2OBI501
Act
54-87
O02
081
12-81
0-18
15-90
12-96
014
008Zone B
B1502
Act
54-09
003
1-43
12-52
023
16-53
12-71
O30
007B1503
Act
53-20
004
2-01
12-01
O20
16-80
12-52
043
006C1501
Act
52-69
006
2-82
11-64
017
1718
12-63
052
005C2501
Act
52-03
008
3-72
11-80
015
16-43
12-49
063
007C4301
Act-Hb
51-37
Oil
4-82
11-45
017
16-23
12-83
075
00545-20
021
915
1602
022
12-74
12-45
1-53
027Zone C
C5501
Act-Hb
5102
014
513
1115
022
16-65
12-77
O81
00746-43
025
8-77
14-89
020
1313
12-37
1-41
005C6501
Act-Hb
49-75
018
5-67
1074
024
17-29
12-83
092
00847-61
033
816
13-92
022
13-91
12-51
1-25
007C7301
Act-Hb
5O37
015
5-35
1051
025
17-40
12-84
087
00848-44
041
7O4
12-85
021
14-54
12-57
103
010X
N
>
Total
Si
Ti
Al
Fe
Mn
Mg
Ca
Na
K97-77
7-896
O002
0137
1-542
O022
3-409
1-998
0040
001497-91
7-775
O003
0242
1-506
0028
3-541
1-957
O083
0O1297-27
7-689
O004
0343
1-452
O024
3-619
1-939
O120
O01097-76
7-573
O006
0478
1-399
0021
3-680
1-945
0145
O00997-40 97-78 97-79 97-96 97-50 97-70 97-98 97-82 9719
Atomic proportions, 0 •= 234)
7-513 7-392 6-730
0009 0012 O024
0633 0817 1-606
1-425 1-378 1-995
O018 0021 0028
3-536 3-480 2-827
1-932 1-978 1-986
0177 O209 0442
O012 0009 00527-327
O015
0869
1-339
0027
3-563
1-965
0226
O0126-864
0028
1-529
1-841
0025
2-892
1-960
O403
0-0097-177
O020
0965
1-296
O029
3-717
1-983
0257
O0146-958
0036
1-406
1-701
O027
3O30
1-959
0355
00127-240
0016
0906
1-264
0030
3-727
1-978
0242
00147O93
0045
1-215
1-573
0026
3173
1-972
0292
0019Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
Sample no.
SiOj
TiOj
A12O3
FeO*
MnO
MgO
CaO
Na2O
K2O
Total
Si
Ti
Al
Fe
Mn
Mg
Ca
Na
KCS501
Hb
49-29
058
6-99
11-30
029
15-67
12-45
117
O09
97-83
7112
0063
1189
1-364
0036
3-369
1-925
0328
O017Zone C
C950I
Hb
46-95
055
9-17
13-27
024
13-68
12-45
1-59
O08
97-98
6-853
0061
1-578
1-620
0O30
2-976
1-947
0451
0014C9301
Hb
45-75
053
1110
1413
023
11-85
1218
1 79
013
97-69
6-725
O059
1-924
1-737
0028
2-596
1-918
O511
O025D1501
Hb
46-41
057
8-97
1113
021
16-54
12-53
1-31
004
97-71
6-739
0062
1-536
1-352
O026
3-579
1-949
0368
O007D2S0I
Hb
47-24
061
6-87
16-61
025
12-40
12-24
099
Oil
97-32
Atomic
7042
0069
1-208
2071
0031
2-755
1-956
0287
0021D2502
Hb
47-91
074
6-85
15-21
025
13-51
12-28
1O0
009
97-84D2503
Hb
48-53
067
6-71
15-24
021
12-99
12-43
084
008
97-70
proportions, 0 = 230
7O50
0082
1188
1-872
0031
2-963
1-936
0285
00187136
O074
1163
1-874
0027
2-847
1-958
0240
O014Zone
D350I
Hb
47-83
088
7-49
14-82
021
12-58
12-48
094
003
97-26
7058
O098
1-303
1-829
0027
2-767
1-973
O270
O005D
D3502
Hb
46-89
095
7-72
14-53
022
13-36
12-44
1O0
004
9715
6-939
O106
1-347
1-798
0028
2-946
1-972
0286
O007D3301
Hb
46O0
083
9-91
13-81
024
13-38
1208
1-46
010
97-81
6-743
O092
1-713
1-693
O030
2-923
1-898
0416
0019D430I
Hb
44-30
1-20
1096
12-93
019
13-39
12O0
201
O04
9702
6-551
0133
1-910
1-599
0024
2-951
1-901
0576
0O07D4302
Hb
44-39
1-47
11O0
11-62
025
14-30
11-94
2-23
0-06
97-26
6-517
0162
1-904
1-426
O031
3129
1-878
0635
0011D4401
Hb
43-80
1-79
11-83
1023
016
1510
11-89
2-36
005
97-21
6-398
0197
2037
1-250
O020
3-287
1-861
0-669
O009H
X
m
X
oo
>z—1
o"0
oH
tn
FeO* means total iron as FeO.
Abbreviations: Act = actinolite, Hb = hornblende.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
410 H. ISHIZUKA
1.0
0.5
0.5
a2
•iu.B -
• :
o:ZONE
ZONE
ZONEB
C
0AMPHIBOLE
I I I I | I I I I | I I I I | I I I. I
0.5 1.0 1.5 2.0
FIG. 7. Variations of (Na + K) and Fe*/(Fe* + Mg) against A1IV for Horokanai amphiboles. Tie-lines represent
coexisting actinolite and hornblende pairs; both analyses are listed in Table 7.
may be due to lack of buffered assemblages in the analysed samples except for sample D1501
which includes chlorite.
In the medium- to higher-grade part of Zone D where pyroxene occurs, the A1IV content of
hornblende increases again with increasing metamorphic grade to approach pargasite.
All the analysed amphiboles are poor in Na(M4) less than 0-14 and well within the
low-pressure field on the A1VI vs. Si diagram of Raase (1974). Furthermore, the hornblendes
coexisting with sphene or ilmenite show a systematic enrichment in TiO2 with increasing
metamorphic grade, ranging from 0-21 wt. per cent (no. C4301) to 1-79 wt. per cent (no.
D4401). This is in harmony with Raase's (1974) empirical rule of temperature dependence of
the TiO2 content.
Pyroxene (Table 8)
Pyroxene develops in the medium- to high-grade part of Zone D, and in the ultramafic
rocks; the compositional variation in the pyroxene quadrilateral is illustrated in Fig. 8.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
TABLE 8
Representative analyses of pyroxenes from Horokanai Zone D and ultramafic rocks
Sample no.
SiOj
TiOj
AJ2O3
Cr2O3
FeO*
MnO
MgO
CaO
Na2O
Total
Si
Ti
Al
Cr
Fe
Mn
Mg
Ca
NaD350I
Cpx
52-01
008
114
003
9-78
0-37
12-45
23-80
032
99-98
1-962
0-002
0O51
0O01
0-309
0012
O700
0962
0024D3502
Cpx
52-08
O07
089
003
9-58
035
1312
23-57
026
99-95
1-962
O002
O040
O001
O302
O011
0737
0952
0019D330I
Cpx
5212
007
117
002
9-12
039
13-75
22-99
024
99-87
1-957
0-002
0052
0001
0286
0012
O770
0925
O018Zone D
D4301
Cpx-Opx
51-95
017
204
002
8-32
029
1413
22-65
O37
99-94
1-940
OO05
O090
0O01
O260
0009
0786
0906
002752-12
004
1-24
002
23-26
036
2213
069
001
99-87
Atomic
1-954
O001
0055
0O01
0729
0011
1-237
0028
0001D4302
Cpx-Opx
52-81
015
1-53
004
7-72
020
14-75
22-43
034
99-97
proportions,
1-961
0004
0067
OO01
O240
0006
0816
0893
002552-73
004
099
002
21-37
035
23-84
062
001
99-97
,0 = 6-0
1-957
0001
0043
0001
0663
0O11
1-319
0025
O001D4401
Cpx-Opx
53-31
016
106
003
714
034
1515
22-26
040
99-85
1-976
O004
0O46
0O01
0221
0011
0837
0884
002953-95
004
056
002
19-28
035
2510
068
001
99-99
1-979
0O01
0024
0O01
0592
0011
1-372
0027
0001Ultramafic
U3010
Cpx-Opx
53-82
001
087
021
2-26
Oil
1905
2315
009
99-57
1-960
OOOO
0O37
0006
0069
0004
1034
0904
O00756-72
001
061
010
702
022
34-55
056
003
99-82
1-969
OOOO
0025
O003
0204
0006
1-787
0021
0002rocks
U302I
Cpx-Opx
53-92
001
084
024
2-20
010
19-16
23-20
010
99-77
1-960
OOOO
0036
O007
0O67
OO03
1038
0904
0O0756-82
001
051
010
6-52
022
34-79
076
003
99-76
1-971
OOOO
0021
OO03
0189
O006
1-798
0028
0O02H
Xm
X
O
70
O
>
>
oX
o
FeO* means total iron as FeO.
Abbreviations: Cpx = clinopyroxene, Opx = orthopyroxene.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
412 H. ISHIZUKA
-Hd
-Fs
75 70 65
FIG. 8. Compositions of Horokanai pyroxenes plotted in the pyroxene quadrilateral. Tie-lines represent
coexisting clino- and orthopyroxene pairs; both analyses are listed in Table 8.
In the orthopyroxene-free metabasites of Zone D, clinopyroxene is salite with Al2O3 of
0-8-1-2 wt. per cent, and shows a systematic decrease of CaO with increasing metamorphic
grade as in the amphibolite facies rocks of Central Abukuma (Shido, 1958) and Broken Hill
(Binns, 1962). These clinopyroxene compositions differ from those of relict clinopyroxenes
persisting in Zone D as porphyroclasts; these relict clinopyroxenes are augites with A12O3 of
4-5-51 wt. per cent.
In the orthopyroxene-bearing metabasites of Zone D, clinopyroxene is salite to augite with
A12O3 of 1-0-2-1 wt. per cent, and orthopyroxene is hypersthene with CaO less than 0-7 wt.
per cent and A12O3 of 0-5-1-3 wt. per cent. The partition coefficient KA = (Fe'/Mg)01"/
(Fe*/Mg)Cpi between coexisting pyroxenes ranges from 1-6 to 1 -8, which is within the limits of
metamorphic pyroxene pairs (K^ = 1-5-2-0; Kretz, 1963). The Mori-Green (1978) thermo-
meter gives equilibration temperatures of 660 °C±40 °C for such K^ values, if the Fe2O3
content is assumed to be negligible.
In the ultramafic rocks, clinopyroxene is diopside with A12O3 less than 0-9 wt. per cent, and
orthopyroxene is bronzite to enstatite with CaO and A12O3 less than 0-8 and 0-6 wL per cent,
respectively. The K^ ranges from 1-6 to 1-7, and gives equilibration temperatures, by the
above thermometer, of 680 °C±20 °C.
Olivine and spinel (Table 9)
Olivine and spinel are common constituents in the ultramafic rocks. The forsterite content
of olivine, Fo = Mg/(Mg + Fe*) = 0-91-0-93, is relatively constant without reference to
rock-types. The NiO content of olivine is 0-35 wL per cent on average. Spinel spans a wide
range of composition, although the Fe3+/(Cr + Al + Fe3+) ratio is small regardless of
rock-types; e.g. the Mg/(Mg + Fe2 +) ratio ranges from 0-53 to 0-72 in pyroxenite, from 0-44 to
0-51 in dunite, and from 0-45 to 0-58 in harzburgite, while the Cr/(Cr + Al + Fe3 +) ratio ranges
from 0-10 to 0-63 in pyroxenite, from 0-78 to 0-88 in dunite, and from 0-67 to 0-77 inDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 413
harzburgite. The TiO2 content of spinel is less than 0-3 wt. per cent. The olivine-spinel
thermometer of Roeder et al. (1979) gives equilibration temperatures of 603-717 °C for
pyroxenite, 598-753 °C for dunite, and 598-729 °C for harzburgite.
TABLE 9
Chemical compositions of olivine and spinel, and temperature estimates for Horokanai
ultramqftc rocks
Sample no.
Pyroxenite
P-3010
P-3021
P-4010
Dunite
D-5010
D-5021
D-5031
D-3010
D-3021
Harzbwgite
H-5010
H-5021
H-5032
H-3010
H-3021Olivine
Mg*
0-918
0-918
0-930
0-918
0-918
0926
0-927
0-929
0-921
O923
0-928
0-926
0-931Mg*
0-719
0-626
0-534
0-507
0494
0-503
0-440
0-461
0-453
0-474
0-584
0528
0523Spinel
Cr**
0096
0355
O628
0832
0853
0847
0878
0781
0766
0714
0711
0671
0681Al**
0804
O580
O338
0089
O130
O060
0064
0178
0230
0254
0251
0300
0290Fe**
O100
0065
O034
0079
0017
0093
0058
0041
O004
O032
0038
0029
0029T***
CQ
717
665
603
753
710
723
630
598
598
603
729
633
610
Mg* = Mg/(Mg + Fe2+); Cr**, Al** and Fe" = Cr/(Cr + Al + Fe3+),
Al/(Cr + Al + Fe3+) and Fe3+/(Cr + Al + Fe3+), respectively. T*** = tem-
peratures by the method of Roeder et al. (1979). Fe2 + of olivine means total iron
asFe2 + ,and Fe3+ and Fe2+ of spinel were calculated from total iron assuming
ideal spinel formula.
METAMORPHIC CONDITIONS
The temperatures of the zeolite fades range from 100 °C to 300 °C (for a review, see Zen &
Thompson, 1974). Boles & Coombs (1977) suggested 190 °C for the first appearance of
iron-rich pumpellyite in the zeolite facies sandstone of New Zealand, and Liou (1979)
estimated temperatures for the laumontite + iron-rich pumpellyite assemblage in the East
Taiwan ophiolite to be at 150-250 °C; the iron content of the Horokanai pumpellyite is
similar to those from New Zealand and Taiwan. Nitsch (1971) demonstrated experimentally
that, at Pfluld = Ptou, = 2 kb, the prehnite + pumpellyite + chlorite + quartz assemblage is
stable up to 345 °C±20 °C, whereas the chemically equivalent assemblage actinolite +
chlorite + epidote + quartz exists above 350 °C; the former assemblage occurs in the
highest-grade part of Zone A, and the latter one is critical to Zone B. The mineral assemblage
transitional from the greenschist to amphibolite facies, i.e. actinolite+calcic plagioclase +
chlorite, is stable within the temperature interval between 475-550 °C at Pnuid = P,^,, = 2 kb
and at the oxygen fugacity of the QFM buffer (Liou et al., 197'4), or between 370-420 °C as
given by the carbon isotope and calcite-dolomite thermometers (Maruyama et al, 1982,
1983). The temperatures of the highest-grade part of Zone D (hornblende-granulite facies)
are 600-700 °C as estimated on the Fe-Mg partitioning of coexisting pyroxenes calibrated byDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
414 H. ISHIZUKA
Mori & Green (1978). The equilibration temperatures of the underlying ultraraafic rocks are
680 °C±20 °C also on the Fe-Mg partitioning of coexisting pyroxenes and 600-750 °C on
the olivine-spinel thermometer of Roeder et al. (1979), similar to those calculated for the
overlying granulite fades rocks of Zone D.
There are few constraints available for the quantitative estimates of the metamorphic
pressures. However, the absence of the typical albite-epidote amphibolite fades assemblage,
the presence of the prehnite + actinolite + chlorite assemblage in the highest-grade part of
Zone A, the presence of the albite + oligoclase + actinolite + hornblende+chlorite + epidote
assemblage in Zone C, and the existence of Ca-amphiboles poor in Al^ as well as Na(M4), all
suggest low-pressure conditions (Shido, 1958; Miyashiro, 1973; Raase, 1974; Brown, 1977;
Maruyama et al, 1983; Liou et al., in press).
Accordingly, it is suggested that the Horokanai ophiolite was progressively metamor-
phosed in the temperature range of 100-750 °C, and its fades series, ranging from the zeolite,
greenschist, amphibolite to granulite fades, belongs to the low-pressure type.
The serpentinization of the Horokanai ultramafic rocks, characterized by the chrysotile +
lizardite assemblage, occurred at temperatures less than 350 °C (Evans et al., 1976; Coleman,
1977). As there is no textural evidence showing recrystallization of serpentine minerals, the
stage of serpentinization may have postdated the equilibration stage of the ultramafic rocks
at the granulite facies. Also, such low-temperature serpentinization may have been unrelated
to the formation of antigorite occurring in one sample of the Horokanai ultramafic rocks
as 'relict' partially replaced by chrysotile or lizardite, because antigorite is stable at
temperatures in excess of 350 °C and as high as 500 °C (Evans et al., 1976; Coleman, 1977). It
is, however, ambiguous to determine the formative stage of 'relict' antigorite.
DISCUSSION
Recently, it has been shown that much of the oceanic crust has suffered metamorphism of
the low-pressure fades series which comprises the zeolite, greenschist or actinolite-caldc
plagioclase, and amphibolite fades (Melson & van Andel, 1966; Cann & Funnell, 1967; Cann,
1969; Ploshko et al, 1970; Miyashiro et al, 1971; Aumento et al, 1971; Bonatti et al, 1975;
Helmstaedt, 1977). Miyashiro et al. (1971) referred to this recrystallization as ocean-floor
metamorphism. However, the structural and stratigraphical relationships among the
dredged metamorphic rocks from the ocean-floor are not known, and they are generally
inferred from similar ophiolites found in Cyprus (Gass & Smewing, 1973), Oman (Coleman,
1977), Newfoundland (Coish, 1977), Chile (Stern & Elthon, 1979), Taiwan (Liou & Ernst,
1979; Liou, 1979), and the California Coast Ranges (Evarts & Schiflman, 1983). The
metamorphic nature of the Horokanai ophiolite is broadly similar to that inferred for
ocean-floor metamorphism, but differs from it, most importantly, in the presence of the
granulite fades rocks (the highest-grade part of Zone D) in the former and their absence in the
latter.
In most of the ophiolites referred to above, the metamorphic grade increases downwards
from the zeolite to low amphibolite fades within a short vertical sequence as it does in the
Horokanai Zone A to the lower-grade part of Zone D, but the maximum metamorphic grade
attained is in an upper part of the gabbro members, below which the metamorphic effect dies
off rapidly, preserving unmetamorphosed gabbros. The orthopyroxene-bearing rocks from
the lower part of the mafic members in these ophiolites have been regarded as magmatic
(commonly called norite or two-pyroxene gabbro), even though few chemical data are
available on their constituents. These observations led Coleman (1977) to deduce that
the grade of ocean-floor metamorphism is less than the low amphibolite fades, andDownloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
THE HOROKANAI OPHIOLITE 415
metamorphism does not take place below the upper part of the oceanic layer 3 (gabbroic
rocks).
However, there is little doubt that the lowermost horizon of the Horokanai gabbro
member (i.e. the lowermost horizon of Zone D) was subjected to granulite fades
metamorphism. A granoblastic texture, pleochroic orthopyroxene, and equilibration
temperatures given by pyroxene thermometers, all strongly support this view. The
underlying ultramafic rocks of the Horokanai ophiolite also record the equilibration
temperatures of the granulite fades as obtained by pyroxene and olivine-spinel thermometers.
Coleman (1977) considered that the high-grade rocks such as the high amphibolite to
granulite fades rocks were part of the basement rocks of island arc ophiolites within the high
heat flow regime such as Yap Islands. The following facts, however, argue against an island
arc origin of the high-grade rocks: (1) the high-grade rocks derived from the gabbroic rocks
actually form an integral part of the ocean-floor (Mid-Atlantic Ridge: Miyashiro et al., 1971;
FAMOUS region: Helmstaedt, 1977); (2) the Horokanai ophiolite has a geochemical affinity
with normal abyssal tholeiites (Ishizuka, 1981).
Miyashiro et al. (1971) and Mevel et al. (1978) offered alternative hypotheses that the
oceanic realms are the site for formation of the high-grade rocks, but the details between them
differ. Miyashiro et al. (1971) pointed out that the high-grade rocks with gneissose or banded
texture are commonly assodated with serpentinites. Because such an association is similar to
that of serpentinized peridotite and contact metamorphic rocks in the Lizard area, they
considered that the high-grade rocks were formed by contact metamorphism by high-
temperature peridotite intrusion in oceanic crust. Mevel et al. (1978) claimed that the
high-grade metamorphism in the Chenaillet ophiolite massif was accompanied by shearing
or sub-horizontal plastic flow. They then suggested that the long-duration of the large
magma chamber near the ridge axis along with the assodated shearing or plastic flow
provides the deep thermal structure responsible for formation of the high-grade rocks.
We were unable to determine which of these hypotheses can best explain the origin of the
Horokanai high-grade rocks, but the following observations are relevant; (1) there is no
distinct gap of thermal structure between Zones C and D, and between Zone D and the
underlying ultramafic rocks, suggesting that the metamorphism progrades continuously
downwards to the ultramafic rocks; (2) the constituent minerals in the high-grade rocks are
homogeneous in composition, and they are completely equilibrated to the granulite facies; (3)
strong deformation or shearing is rare in the granulite zone; and (4) there is no intrusion of
ultramafic rocks within the granulite zone. It is therefore concluded that the metamorphism
of the Horokanai high-grade rocks prevailed under relatively static and high-temperature
conditions, and had enough time to give rise to chemical equilibrium. These facts suggest that
the granulite facies conditions were attained in part of the oceanic lower crust and upper
mantle. In light of these observations, the concept of ocean-floor metamorphism requires
further consideration.
ACKNOWLEDGEMENTS
I wish to express my sincere thanks to Professor S. Banno for his critical reading of this
manuscript as well as for his never-ending encouragement throughout this study, and to
Messrs N. Gouchi and M. Imaizumi with whom I enjoyed frank discussions. I am also
indebted to Professor T. Suzuki for his valuable suggestions and continuous encouragement.
I sincerely thank B. W. Evans, J. G. Liou, and W. E. Trzdenski, Jr. whose careful reviews
greatly improved the final version of the manuscript. This study was financed in part by
Grant-in-Aid for Sdentific Research of Ministry of Education of Japan (No. 56740351/
Ishizuka).Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
416 H. ISHIZUKA
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island arc. Earth planet. Sci. Lett. 13, 167-74.
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their implications for ocean floor metamorphism. Tectonophysics, 55, 179-213.
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British Columbia. Bull. geol. Soc. Am. 84, 1911-22.
Tagiri, M, 1973. Metamorphic rocks of the Hitachi district in the southern Abukuma Plateau, Japan. Sci. Rep.
Tohoku Univ. Ser. Ill, 12, 1-67.
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Plateau, Japan. Contr. Miner. Petrol. 62, 271-81.
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Earth planet. Sci. 2, 179-212.Downloaded from https://academic.oup.com/petrology/article/26/2/391/1473978 by College of Law Library user on 07 March 2025
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The Island Arc (1997) 6, 2-24
Thematic Article
Contrasting two types of orogen in Permo-Triassic Japan:
Accretionary versus collisional
YUKIO ISOZAKI
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 0-okayama, Meguro,
152 Tokyo, Japan
Abstract Proto-Japan originated from a continental margin of the Neoproterozoic Yangtze
(South China) craton. It represents a unique Permo-Triassic tectonic setting in western
Panthalassa, where two distinct types of orogenic belt occurred side by side. There was an
accretionary orogen between the Yangtze craton and the Proto-Pacific (Farallon) Plate and a
collisional orogen between the Sino-Korean (North China) and Yangtze cratons. This article
reviews results of the latest on-land geological studies concerning Permo-Triassic tectonics in
Japan and proposes a new plate tectonic interpretation as well as a paleogeographic reconstruc-
tion of this particularly unique geotectonic regime. Special emphases are given to (i) the
accretion processes and products derived by collision-subduction of the Permian Akiyoshi
paleoseamount and Maizuru paleo-oceanic plateau; (ii) the field occurrence of 220-Ma Sangun
high-P/T schists and its implication for the exhumation process and ‘tectonic sandwich’
structure; (iii) the extensive development of a subhorizontal nappe of the pre-Jurassic rocks and
their bearing on the orogenic edifice; and (iv) the restricted occurrence of the 250-Ma collision
complex in the Hida and Oki belts and the relevant connection to the Precambrian cratons and
collision suture in East Asia. The newly proposed paleogeographic reconstruction is also tested
by faunal provinciality of Permo-Triassic fossils from shallow-water sediments.
Key words: accretion, collision, high-PIT schists, Japan, nappe, ophiolite, Permo-Triassic, sea-
mount, Sino-Korea, Yangtze.
INTRODUCTION
The predominance of accretionary complexes (AC)
and the association of detached continental frag-
ments in Japan suggest that the Japanese Islands
have fundamentally developed through conver-
gence between oceanic and continental plates along
active margins. The origin of Japan goes back to
the Neoproterozoic era (ca 750-700 Ma) when the
proto-Pacific basin was formed by the break-up of
the supercontinent Rodinia (Hoffman 1991; Dalziel
1992; Powell et al. 1993; Park et al. 1995). After
the conversion from passive to active continental
margin in the Early Paleozoic, Pacific-ward oro-
genic growth of Japan was achieved successively
by subduction-accretion processes from the Late
Paleozoic to the present (Isozaki & Maruyama
Accepted for publication July 1996 1991; Isozaki 1996). With regard to the evolution
of the Japanese Islands, the Permo-Triassic tecton-
ics are particularly important because the funda-
mental framework of the Mesozoic-Cenozoic oro-
genic belt of Japan was established and stabilized
at that time by virtue of the generation of two
orogenic belts of distinct type side by side; that is,
the subduction-accretion type and the continent-
continent collision type.
The Permo-Triassic tectonics of the Japanese
Islands also provide a vital piece of information for
reconstructing the paleogeography of the western
Panthalassa (proto-Pacific) margin, because this
period corresponds to the zenith of the superconti-
nent Pangea (Klein (ed.) 1994; Embry et al. (eds)
1994). Compared with the Tethyan peripheries (or
paleo-Tethyan [Sengor 1989]), information has
Accretion and collision in Permo-Triassic Japan 3
supercontinent, engulfing a closed ocean basin
called Tethys (or paleo-Tethys). Both the Sino-
Korean and Yangtze Blocks are composed of Pre-
cambrian rocks including Archean cratonic core
(Jahn 1990; Zheng et al. 1991; Zhou 1994), and
are regarded as having detached originally from
the supercontinent Rodinia before 700 Ma. The
polar wander history of these two continental
blocks has been partly documented (Lin et al.
1985; Powell et al. 1993), and at ca 250 Ma they
collided against each other to form a larger conti-
nental piece which nearly corresponds to the major
part of the present China.
The collisional boundary between these two
blocks is known as the Qinling-Dabie suture in
central China marked by a 230-Ma ultrahigh-
pressure metamorphic belt (Wang et al. 1989;
Okay et al. 1989; Maruyama et al. 1994; Cong &
Wang 1995). The eastern location of the suture is
well established up to the Shandong Peninsula,
Northeast China (Yang & Smith 1989; Hirajima et
al. 1990; Enami & Zang 1990), but further recog-
nition to the east remains debatable. Judging from
the similarity in protolith lithology and age, Sohma
et al. (1990) regarded the Hida belt (including the
Unazuki schist belt) in central Japan as an eastern
extension of this collision suture. In addition, fol-
lowing the comment by Ernst et al. (1988) and
Cluzel et al. (1992), Sohma and Kunugiza (1993)
and Isozaki (1996) nominated two alternative can- been limited or almost absent regarding tectonics
as well as the paleoenvironment of the western
proto-Pacific domains.
In this article, I describe the Permo-Triassic
orogenic units of the Japanese Islands and discuss
the nature of two distinct orogenic belts developed
in proto-Japan, highlighting some diagnostic as-
pects, such as the subduction-accretion of the
reef-capped Akiyoshi paleoseamount and the Mai-
zuru paleo-oceanic plateau, exhumation of the
220-Ma Sangun high-P/T meta-AC, the formation
of a large-scale Kurosegawa nappe of pre-Jurassic
rocks, and generation of 250-Ma medium-pressure
type Hida gneiss through continent-continent col-
lision along the western Panthalassa margin.
TECTONIC OUTLINE OF PERMO-TRIASSIC PROTO-JAPAN
A brief summary is given here of the large-scale
tectonic background of the western Panthalassic
margin where proto-Japan was situated during the
Permo-Triassic. Two isolated continental pieces
named the Sino-Korean (or North China) block and
Yangtze (or South China) block were then located
in the eastern margin of the supercontinent Pan-
gea, as shown in the reconstructed paleogeo-
graphic map (Fig. I). Together with other smaller
continental pieces, these two blocks almost entirely
bridged the northern and southern ends of the
Fig. 1 Tectonic framework of the western Panthalassa at the latest Permian time, ca 255 Ma (modified from Scotese & Langford 1995). Note the location of
proto-Japan at the northeastern edge of the Yangtze craton on the Pacific side
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4 Y. Isoxaki
didates for the suture in the Korean Peninsula; the
Imjingan belt in North Korea, or the Ogchon zone
in South Korea.
The Permo-Triassic orogenic complexes occur
mostly in Southwest Japan, extending 2000 km
along the arc from central Japan to the southern
Ryukyu Islands close to Taiwan (Fig. 2). They
always occupy the Pacific side of the putative
location of the YangtzelSino-Korean suture which
is in China and Korea. This suggests that proto-
Japan was attached to the margin of the Yangtze
Block (Isozaki & Maruyama 1991; Isozaki 1996),
rather than the Sino-Korean block, otherwise an
unrealistic large-scale strike-slip fault is needed
with along-arc displacement of more than 2000 km
within Japan. Accordingly, the Japan side of the
Yangtze craton is regarded to have faced the
proto-Pacific, not directly to paleo-Tethys (Fig. 1).
After a collision betvveen the Yangtze and Sino-
Korean blocks, proto-Japan kept growing around
the amalgamated blocks, and started receiving
I
I A$? Ishigaki Island / materials also from the Sino-Korean craton. Other
parts of Asia, such as Siberia, Tarim, and Burya
(Amuria), were scattered along the periphery of
Pangea in the Permo-Triassic period.
PERMO-TRIASSIC OROGENIC UNITS IN JAPAN
Distribution of the Permo-Triassic orogenic units in
Japan is depicted in Fig. 2. It is remarkable that
the occurrence of Permo-Triassic orogenic units is
mostly limited to Southwest Japan and the Ryukyu
Islands. There is no mappable-sized example of the
Permo-Triassic orogenic unit recognized in North-
east Japan and Hokkaido, excepting the Hitachi-
Takanuki Belt at the southern tip of Northeast
Japan. Although minor amounts of Permo-Triassic
sedimentary rock occur also in Northeast Japan,
they represent overlapping continental shelf-slope
sequences that accumulated upon the Early-Middle
Paleozoic complexes (older continental crust and
b B
Japan Sea
Permo- Triassic
Orogenic Units in Japan Taishaku area
500 km
I
Continental block
Permiam AC High-PTT metamorphosed Permian AC (with Triassic age)
Tectonic outliei of Permo-Triassic units (Kurosegawa klippe)
Fig. 2 Distribution of the Permo-Triassic orogenic units in Japan (modified from lsozaki 1996) and localities of referred units and areas mentioned in the text. Ak,
Akiyoshi belt, Mz, Maizuru belt; Sn, Sangun belt; HT, Hitachi-Takanuki belt; SK, Southern Kitakami belt; MTL, Median Tectonic Line; I-KTL, Ishigaki-Kuga Tectonic
Line; TTL, Tanakura Tectonic Line.
14401738, 1997, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1997.tb00038.x by Susam Welch - Ohio State University University Libraries , Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Accretion and collision in Permo-Triassic Japan 5
southwestern extension of the Permo-Triassic ac-
cretionary orogen in Southwest Japan. There is no
known occurrence of the collision-related Permo-
Triassic unit to date either in the Outer Zone of
Southwest Japan or in the Ryukyus.
The Permo-Triassic AC in the Inner Zone of
Southwest Japan comprises three units distributed
in three distinct belts; the Akiyoshi, Maizuru and
Sangun belts (Fig. 2). The AC in the Sangun belt
was metamorphosed at high-P/T blueschist facies.
As demonstrated schematically in Fig. 3, all three
units occur as subhorizontal nappes, forming large-
scale imbricated bodies. These units are sand-
wiched tectonically between the older and over-
lying 300-Ma blueschists of the Renge belt and the
younger and underlying Jurassic AC of the Mino-
Tanba belt, respectively (Hayasaka 1985; Isozaki
& Itaya 1991; Kabashima et al. 1993). The basal
boundary fault (called the Ishigaki-Kuga tectonic
line) can be traced for -2000 km along Southwest
Japan and the Ryukyus (Isozaki & Nishimura
1989).
The non-accretionary units occur in the Hida and accretionary complexes) or olistostromal fragments
incorporated in younger Jurassic-Cretaceous ac-
cretionary complexes. These units are not dealt
with here because of their lesser tectonic signifi-
cance to the present study.
Southwest Japan is subdivided into the Inner
Zone, on the Japan Sea side, and Outer Zone, on
the Pacific side, and the boundary is presently
demarcated by a distinct fault called the Median
Tectonic Line (MTL). As shown in Fig. 2, Permo-
Triassic orogenic complexes have a clear double-
belted distribution pattern; namely, one large area
on the Japan Sea side of the Inner Zone and the
other discontinuous narrow belt in the middle of
the Outer Zone. The major components of the
Permo-Triassic orogen in Southwest Japan are, on
one hand, ancient accretionary complexes (AC) and
their high-P/T metamorphosed equivalents that
were generated by oceanic subduction from the
Pacific side, and on the other hand, continent-
continent collision-related orogenic units. In the
Ryukyu Islands, high-P/T metamorphosed AC on
Ishigaki Island, close to Taiwan, represent the
+- Pre-Jurassic complexes -
of the Inner Zone piGa]
I Kuroseaawa I
Tectonic Outlier I (Klipie, I
Nagato-Hida
Marginal T.L.
I S N
,I
Crel -Paleog Shimanto complexes
LA INNER ZONE - - - ourm ZONE
1 Mid-Cretaceous]
Pre- Jurassic
Pre-Japan Sea complexes
N voIc. (Aklyoshi Orogen) S
I Jurassic complex -
fi -- I ASakawa Orogen)
- .a_--
high-PIT met
./-----:, '/A/ '
1' / I==> ,' / oceanward orogenic growth i
v 9
oo2
P
Fig. 3 Fundamental structure of Southwest Japan (Isozaki & ltaya 1991) (a) Schematized geotectonic profile of Southwest Japan at present Ch, Chichibu belt,
Ry, Ryoke belt (b) Schematized profile of Cretaceous Southwest Japan prior to the klippe formation, (c) Schematized map view of Cretaceous Southwest Japan
14401738, 1997, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1997.tb00038.x by Susam Welch - Ohio State University University Libraries , Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
6 Y. Isoxaki
Oki belts in Southwest Japan, and these are com-
posed of polymetamorphosed gneisdschist com-
plexes with a signature of 230-Ma regional meta-
morphism of the medium-pressure type facies
series (Komatsu 1990). In the Outer Zone, a nar-
row domain called the Kurosegawa belt (or tectonic
zone, terrane etc.; Yoshikura et a,l. 1990; Isozaki et
al. 1992) represents a tectonic outlier of the equiv-
alent pre-Jurassic units in the Inner Zone and will
be described later. Mesozoic cover sediments un-
conformably resting on these orogenic rocks are
minor in volume but they provide excellent age
constraints for the Permo-Triassic tectonics. In the
following sections, the characteristics and mode of
occurrence of these Permo-Triassic orogenic units
are briefly described.
PERMO-TRIASSIC ACCRETIONARY COMPLEXES
Permo-Triassic AC in Southwest Japan consist of
three subhorizontal nappe units; the weakly meta-
morphosed Permian AC of the Akiyoshi belt, the
220-Ma high-P/T metamorphosed AC of the San-
gun belt, and the weakly metamorphosed Permian
AC of the Maizuru belt, from top to bottom. The
Akiyoshi AC and the Maizuru AC experienced
regional metamorphism of lower greenschists
facies or less, and yield fossils useful for dating
and for assigning their accretion setting at an
ancient trench, through oceanic plate stratigraphy
(OPS) analysis (Isozaki 1996). Each AC is nowhere
unique. The Akiyoshi AC is characterized by
kilometric-sized reef limestone blocks of Car-
boniferous-Permian age. The Maizuru AC contains
a dismembered ophiolite suite, known as the
Yakuno Complex. The Sangun AC, on the other
hand, was metamorphosed into blueschists with
strong deformation but retain protolith assemblage
suggesting an AC origin.
AKlYOSHl AC
The Permian Akiyoshi AC occurs as a subhorizon-
tal veneer-like nappe of approximately 2000 m
thick, occupying the highest structural horizon
among the Permo-Triassic orogen, above the San-
gun belt (Fig. 3). The Akiyoshi AC consists of
rocks derived both from continental and oceanic
crusts; that is, terrigenous clastics such as gray-
wacke sandstone, mudstone, conglomerate, and
acidic tuff, as well as bedded radiolarian and/or
spicular chert, basaltic greenstones, and reef lime-
stone. Most parts are occupied by chaotically mixed units characterized by a block-in-matrix texture
(Kanmera & Nishi 1983). Some parts are repre-
sented by a series of tectonically imbricated thrust
sheets composed mainly of coherent terrigenous
clastics (Hara & Kiminami 1989).
One of the best examples of this AC occurs in
the Akiyoshi area (Fig. 4), where large slabs or
blocks of Carboniferous-mid-Permian limestone
are enveloped in a matrix of coarse-grained elastic
rocks of Late Permian age. The limestone blocks
vary from pebble-sized pieces to the biggest block
of -3 x 5 km. Field occurrence and fossil ages
indicate that the older limestone blocks are allo-
chthonous clasts secondarily incorporated into the
argillaceous matrix of younger age.
Bedded cherts and limestone in the Akiyoshi AC
are regarded as ancient pelagic sediments and
shallow-water organic reef complexes developed
0 terrigenous clastics chert
a limestone (isolated small blocks) limestone (large block) Ev-J basaltic greenstones
7 limestone clastics
Y
I
CI-7, P1-4: Carboniferous and Permian fossil zone
Fig. 4 Geologic map of the Akiyoshi area, Southwest Japan (a, simplified
from Kanmera & Nishi 1983) and schematic cross-section of a large limestone
body (b, modified from Sano & Kanmera 1991). The Permian Akiyoshi AC
consists of the Late Permian terrigenous clastics with numerous allochtho-
nous blocks of the Carboniferous-middle Permian limestones and cherts plus
basaltic greenstones. (a) Note the scattered limestone blocks within the
younger terrigenous clastics in the southwestern part of the area. (b) Appar-
ently large and thick limestone bodies are, in fact, comprised of numerous
smaller fragments that occurred in stratigraphically random order.
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Accretion and collision in Permo-Triassic Japan 7
1979). Litho- and biofacies analyses of the lime-
stones indicate a shallow-water reef environment
for their origin (Ota 1968). In particular, a capping
reef on top of a basaltic seamount that is isolated
from continental influence is most likely because of
the absence of coarse-grained terrigenous clastics.
Lithology, texture and field occurrence indicate
that these rocks represent remnants of a collapsed
ancient mid-oceanic seamount complex (Kanmera
& Nishi 1983; Sano & Kanmera 1991). Faunas
from the Akiyoshi limestone are generally of
Tethyan type, however, Early Carboniferous corals
show a strong affinity with those found in Australia
(Kato 1990), suggesting its potential proximity to
Gondwana.
In contrast, bedded chert from this AC repre-
sents a sedimentary facies of deep-water origin
(Kanmera & Nishi 1983; Uchiyama et al. 1986;
Goto 1988). The age of the bedded chert almost
completely overlaps with that of the reef limestone
(Fig. 5), indicating that these two distinct oceanic
lithologies are regarded as lateral equivalents in
the same ocean basin. The limestone-greenstone
complex represents rocks of an ancient seamount,
while the cherts represent sediments accumulated
in the deep sea below the carbonate compensation
depth (CCD) around the seamount (Fig. 6). Older
limestone clasts are sporadically contained within
younger cherts, and in addition, allodapic limestone
(calcareous turbidite) are interbedded with these upon basaltic seamounts, respectively. A tectonic
setting of a convergent plate boundary between a
continent and an ocean, namely, a trench, is indi-
cated by this lithologic assemblage and field occur-
rence (Kanmera & Nishi 1983; Kanmera et al.
1990). Using microfossils such as fusulinids, cono-
donts and radiolarians, a detailed age determina-
tion of the primary OPS has been reconstructed
(Fig. 5). This stratigraphic reconstruction indicates
that the Akiyoshi AC was formed at ca 260 Ma by
the subduction of -80-million-year-old oceanic
plate (probably the Farallon Plate according to
Engebretson et al. 1985 and Maruyama & Sen0
1986) along a continental margin of Yangtze. The
Akiyoshi AC is unconformably covered by Late
Triassic and Early Jurassic shallow-water clastics.
COLLAPSE OF THE AKlYOSHl PALEOSEAMOUNT
Compared with other AC units in Japan, the Aki-
yoshi AC is unique in having huge reef limestone
blocks of mid-Carboniferous-Permian age that are
associated with basaltic greenstones (Fig. 4). The
fusulinid and rugose coral biostratigraphy indicates
that the primary thickness of the Akiyoshi lime-
stone was -700 m (Toriyania 1958; Hase et al.
1974). The base of the limestone (Visean) rests
conformably upon pillowed alkali basalts (less than
100 m thick), which are distinct from typical mid-
ocean ridge basalt (MORB) (Hase & Nishimura
260 Mi
Fig. 5 Oceanic plate stratigraphy (OPS) of the 295 M:
Permian Akiyoshi accretionary complex (AC)
demonstrated in three isolated areas (Yasuba-
Shirakidani, Taishaku and Akiyoshi areas) in
Southwest Japan (Isozaki 1987) h, OPS for sea-
mount (topographic high), d, OPS for deep-sea
floor, m, OPS for marginal flank of seamount
Note the coeval development of three distinct
facies as lateral equivalents -aii columns not to scale
,d,~m, h U I
,d, siliceous mudstone Is conglomerate
4, a acidic tuff limestone I,
chert greenstones
\I:#,! ,I
4 YASUBA~SHlRAKlDANl 3 TAISHAKU :d!
14401738, 1997, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1997.tb00038.x by Susam Welch - Ohio State University University Libraries , Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
8 Y. Isozaki
reef limestone + OIB
into AC as
lensoids/blocks '
I I I ^._ I
deep-sea
pelagic chert
?"w.&am ..
vyvvvvvvvv "'Bl..2
dv v v v v v v v v v v ' paleo-seamount
vvvvvvvMORB vvvv not to scale - deep-sea facies I u. marginal facies- lopo-high facres marginal facies - - -
Fig. 6 Topographic reconstruction of the paleo-seamount with a reef caD and deep-sea floor (modified from lsozaki 1987) Refer to Fig 5 for the OPS of three
distinct facies
cherts (Hase et al. 1974; Isozaki 1987). These
indicate a lateral lithofacies change between
shallow-water limestone and coeval deep-water
chert, according to paleotopographic relief maps.
The chaotic occurrence in the field of these
limestone-greenstone complexes within terrige-
nous clmtics (Fig. 4) indicates a collapse of the
primary topographic relief of the seamount prior to
the subduction (Sano & Kanmera 1991). Figure 7
illustrates a possible process for seamount collapse
and successive accretion at the trench. An exten-
sional tectonic regime usually appears at the hinge
of the subducting plate just in front of a trench,
triggering the gravitational collapse that is related
to normal faulting. The accreted reef limestone in
the Akiyoshi AC is a fault-bounded lensoid body in
which ill-sorted, debris-flow-like megabreccia oc-
curs. A modern analog for the collapse and accre-
tion of oceanic seamounts can be observed in the
present western Pacific, off Northeast Japan, where
the Daiichi-Kashima and Erimo seamounts are
currently entering the active trench (Cadet et nl.
1987).
THE MAIZURU AC: A PALEO-OCEANIC PLATEAU ACCRETION
The Permian Maizuru AC occupies the lowest
structural level in the Permo-Triassic orogenic
units (Hayasaka 1987). Fossils from this unit are
not so abundant as those in the Akiyoshi AC, but
OPS analysis indicates that the Maizuru AC was
formed in the Late Permian, more or less the same
time as the Akiyoshi AC. The main difference
between the Maizuru and the Akiyoshi AC lies in
lithologic assemblage: (i) the occurrence of 280-Ma
ophiolitic rocks; (ii) the absence of chert; and (iii)
the presence of minor amounts of limestone. The
Maizuru AC in this article includes a unit some-
times called the Ultra-Tanba belt (Ut in Fig. 3a;
Caridroit et al. 1982; Ishiga 1990).
The ophiolitic rocks in the Maizuru AC are
-8 km thick, and are collectively called the Yakuno 2
3
4
'-1
5
* .'
mcollapse products mredeposltcd collapse products
Fig. 7 A model for collapsing paleoseamount at trench (Sano & Kanmera
1991). Note the prE-accretion brecciation of limestone due to the gravitational
collapse of the seamount at the trench.
Complex. They include basalts, gabbros, ultramafic
rocks, and minor amounts of cover sediments and
felsic intrusives (the Maizuru granite). These rocks
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Accretion and collision in Permo-Triassic Japan 9
1990). Under the circumstances a multiple origin
of the Yakuno complex was proposed that assumes
a complicated paleogeographic setting including
arc, hotspot and spreading ridge features in a small
area just like the present southwest Pacific.
However, metamorphic petrology of the suite
provides an important constraint on its origin. The
pyroxenite member near the mafic/ultramafic
boundary (regarded as an ancient Moho) in the
suite reached the granulite facies, suggesting a
high pressure between 5-10 kbar. This requires
unusually thick (15-30 km) oceanic crust, which is
nearly three to six times the average thickness of
modern examples (Ishiwatari 1985). In addition,
scarcity of capping carbonates indicates the crust
surface was beneath the CCD at that time. Given
the depth of the CCD of -3000 m deep, this
suggests that the topographic relief from the deep-
sea floor was less than 2 km. On the basis of these
observations, presented here is another possible
explanation for the origin of the Yakuno Complex.
Oceanic plateaus, such as the Ontong-Java, are
the biggest topographic reliefs in the oceanic do-
main except for the mid-oceanic ridge, although its
internal structure has not yet been revealed in
detail. Judging from their external configuration,
however, an extraordinarily thick oceanic crust,
that is, more than 15 km as estimated for the
Yakuno Complex, can be expected for plateaus.
Abnormally thick oceanic plateau crust appears
favorable to compensate the extra-overburden of a
huge relief of up to 2 km high. The Yakuno ophio-
lite complex probably represents a fragment of
a paleo-oceanic plateau located on the Farallon
Plate. The initiation of a plume is suggested as
giving rise to an oceanic plateau (Maruyama
unpubl. data).
On the other hand, the dismembered feature of
the ophiolite was secondarily given through the
accretion process of an oceanic plateau at a sub-
duction zone. Telescoping of thick oceanic crust
into a dismembered ophiolite suite was not likely to
be achieved by the commonly accepted ‘obduction’
process for ophiolite but probably by subduction-
related underplating, because the structural hori-
zon of the Yakuno Complex beneath the Akiyoshi
AC is consistent with the downward-younging
growth polarity in piled AC nappes in Southwest
Japan (Isozaki 1996). are sliced and chaotically mixed in part, not show-
ing the typical Troodos-style ophiolite stratigraphy
(Ishiwatari 1985; Fig. 8). In particular, the basaltic
unit (-4 km thick) appears abnormally thick in
comparison with other ophiolites. The igneous age
of the basalt-gabbro suite is dated to be ca 280 Ma,
while the age of the sedimentary cover is ca
260 Ma. Basalts of this suite have been analyzed in
terms of petrography, bulk and trace element
chemistry and isotopic composition. However, re-
sults cannot pinpoint any unique tectonic setting
because petrochemical signatures suggesting af-
finity to a mid-oceanic ridge, arc, back-arc, and/or
hotspot were obtained at the same time from
various parts of the ophiolite (Ishiwatari et al.
Ophiolite Lithology Metamorphism
(km) Succession
+mudstone _r
l~ &
basalt
Y lli 3
k. u
u
C
J metabasall/ +
e, Qz-diorite .-
arnphibolite
transition
rnetagabbro
Y
paleo-Moho 3 prehnite-pumpellyite f.
(greenvchistf.)
epidote-amphiboliteJ
amphiholite f.
+---
granulite f.
Fig. 8 Reconstructed ophiolite stratigraphy of the Yakuno complex in the
Permian Maizuru AC, Southwest Japan (simplified from Ishiwatari 1985) This
complex represents a dismembered ophiolite, however the mafic/ultramafic
transition horizon IS correlated with the paleo-Moho surface on the basis of
the metamorphic aspects This suggests that the original thickness of the
oceanic crust may have attained more than 15 km 220-MA HIGH-P/T METAMORPHOSED SANGUN AC
Occupying the middle structural horizon of the
Permo-Triassic orogenic complex in the Inner
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10 Y. Isoxaki
Zone, the Sangun AC occurs as a subhorizontal
nappe beneath the Akiyoshi AC nappe and above
the Maizuru AC nappe (Nishimura et al. 1989;
Isozaki & Maruyama 1991; Fig. 3). Its along-arc
extent is -1500 km in Southwest Japan and the
Ryukyu Islands, and its thickness is less than
2 km. This unit is composed mostly of pelitic to
mafic schists with minor amounts of siliceous and
psammitic schists and rarely with metagabbro and
marble. The radiometric ages of the Sangun
meta-AC concentrate in the 230-200-Ma range,
and its metamorphic peak is suggested to be ca
220 Ma as established by Rb-Sr whole-rock and
K-Ar white-mica measurements (Nishimura 1990).
The grade of the regional metamorphism reaches
the greenschist to glaucophane schist facies, sug-
gesting subduction-related high-P/T metamor-
phism in a Triassic subduction zone with a maxi-
mum pressure of -5-6 kbar. Due to its strong
recrystalization, OPS analysis using microfossils is
not available for this unit. Lithologic associations,
however, indicate that this unit represents an AC
formed at an ancient trench prior to the high-P/T
metamorphism, and thus the age of accretion must
be older than 220 Ma. Nishimura et al. (1989)
reported Middle-Late Permian (ca 260 Ma) micro-
fossils from a chlorite-pumpellyite-bearing phyllite
unit, which has a 220-Ma K-Ar age for metamor-
phic white mica. This unit is closely associated with
the Sangun blueschists that include 225-240-Ma
metagabbroic blocks. On the basis of these rela-
tions, the protolith age of blueschists, namely the
pre-metamorphic accretion age of the main high-
PIT Sangun meta-AC at the trench, is regarded as
being ca 250-230 Ma.
The Sangun meta-AC is tectonically sandwiched
between the over- and underlying nappes of non-
high-P/T AC by sharp fault planes (Fig. 3). The
pressure difference across the faults between the
Sangun meta-AC and adjacent non-high-PIT AC
suggest an -3-4-kb difference corresponding to an
overburden of -9-12 km of crustal material. This
relation imposes the following two constraints, that
is, (i) crustal materials nearly 10 km thick must
have been removed to juxtapose the Sangun meta-
AC and the overlying Akiyoshi AC; and (ii) the
Sangun meta-AC must have gone across the same
crustal thickness to overlie the Maizuru AC. Devel-
opment of a normal fault with displacement of - 10 km across-crust is therefore suggested along
the upper surface of the Sangun meta-AC, while a
reverse fault of the same order of magnitude along
the lower surface. Accordingly, from the primary
high-PIT metamorphic domain along the Benioff zone at -20 km deep to the near-surface domain, a
thin nappe of the Sangun meta-AC was tectoni-
cally uplifted (or exhumated) by activation of the
paired faults of opposite nature, as suggested by
Maruyama (1990). As the nappe of the Sangun
meta-AC occurs extensively along arc for 1500 km,
the mechanism of the exhumation was not of a
local extent but regional, as well as the high-P/T
metamorphism per se.
The timing of the final surface exposure is poorly
constrained. However, because the preservation of
blueschists requires a quick exhumation from the
deep metamorphic domain to avoid thermal anneal-
ing, the emplacement of the Sangun meta-AC into
a shallower crustal level must have occurred imme-
diately after the peak metamorphism (220 Ma),
thus probably in the Late Triassic-Early Jurassic.
The oldest evidence for their surface erosion is the
presence of 200-Ma schist clasts in the Late Cre-
taceous (87-83 Ma) intra-arc sediments, marking
the youngest age limit for the exhumation (Isozaki
& Itaya 1989).
On the other hand, the emplacement of the
Sangun schist nappe over the weakly metamor-
phosed Jurassic AC (160 Ma) by thrusting con-
strains the exhumation to shallow crustal levels
and to having occurred no earlier than 160 Ma. In
addition, the regional intrusion of 85-Ma I-type
granitoids into the Sangun nappe also roughly
constrains the timing of emplacement into shallow
crustal levels. A tectono-metamorphic history for
the Sangun meta-AC is summarized in Fig. 9,
showing a possible path in burial depth versus time
for the Sangun meta-AC unit. This trajectory is
unique to the Sangun unit, and quite distinct from
those of the older 400-300-Ma Renge high-P/T
unit and the younger 100-Ma Sanbagawa high-P/T
unit in Southwest Japan (Isozaki 1996).
KUROSEGAWA KLIPPE: A TECTONIC OUTLIER OF
PRE-JURASSIC COMPLEXES
The Kurosegawa belt in the Outer Zone of South-
west Japan has long been a key geotectonic unit
for understanding the orogenic structure of the
Japanese Islands. This belt features a great variety
of rock types including the Permian AC, the
220-Ma high-P/T schists, fragments of 400-Ma
granite-gneisses, the 400-300-Ma high-PIT
schists, a serpentinite, plus weakly to non-metam-
orphosed Siluro-Devonian and Mesozoic sediments
of continental shelf origin (Maruyama et al. 1984;
Yoshikura et al. 1990; Isozaki et al. 1992). These
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Accretion and collision in Permo-Triassic Japan 11
’ Age carbon! Permian Triassic Jurassic Cretaceous
(Ma) ’ 3y0 250 200 150 100
IIII II~IIII,,,
20 -
25 -
(km)
Tectonic
various components iusually occur as lensoid bodies,
and range in size from kilometers to meters. They
are randomly distribluted and commonly separated
from each other by serpentinite. Given these fea-
tures, the belt has been often described as a
serpentinite mklange zone. In contrast to such
variety in rock types, the distribution is highly
restricted to an extremely narrow zone (Fig. 2),
usually less than 3 km in width, and the belt is
discontinuous along its length.
As the Outer Zone of Southwest Japan is domi-
nated by Jurasso-Cretaceous to Tertiary AC and
their metamorphic equivalents, the pre-Jurassic
orogenic complexes of the Kurosegawa belt appear
as a unique area in the Outer Zone. These pre-
Jurassic units are isolated from the main distribu-
tion of the those in the Inner Zone (Fig. 2). Various
plate tectonic models and interpretations for the
origin of the Kurosegawa belt have been proposed;
for example, a remnant of an ancient Benioff zone,
a rifted continental margin, a collided microconti-
nent or arc, a strike-slip fault zone etc. (Maruyama
et al. 1984; Taira & Tashiro 1987). None of them,
however, can explain all aspects of this belt. A
klippe model was recently proposed for the Kuro-
segawa belt, in which the belt is regarded as a
tectonic outlier of the pre-Jurassic rocks in the
Inner Zone (Isozaki & Itaya 1991; Fig. 3) on the
basis of following lines of evidence.
First, almost all of these sedimentary, igneous,
and metamorphic components of the Kurosegawa
Belt can be identified and correlated with those in
the pre-Jurassic be1t.s of the Inner Zone in terms of sssz
peak metamorphism
(220 Ma)
max. pressure: 5-6 !d~
I I I
I surface of arc crust ocean floor trench 1 Benioff zone I fore-arc prism/crust I
(open ocean) (near continent) I I i
I I
lithology, age, and even faunal provinciality. In
addition, the timing of regional tectonics including
high-P/T metamorphism, formation of an accre-
tionary complex and major unconformity-forming
tectono-sedimentation are identical strictly within
the Inner Zone/Kurosegawa belt and nowhere else
in East Asia. Correlative components and events
of pre-Jurassic belts in the Inner Zone and the
Kurosegawa belt are listed in Table 1. Age and
lithology of components, tectono-sedimentary
events, and faunal provinciality all support the
consanguinity of the pre-Jurassic orogenic units
presently distributed in two separated zones in
Southwest Japan. Concerning the pre-Cretaceous
tectonics in Southwest Japan, the previously ac-
cepted distinction of the Inner and Outer zones is
unnecessary.
Second, all these pre-Jurassic rocks overlie the
Jurassic AC, both in the Inner and Outer Zones. A
remarkable example of a klippe of the Kurosegawa
rocks was first found in central Shikoku (Isozaki &
Itaya 1991; Fig. 10). In this area, 230-Ma high-
PIT schists sit tectonically upon the Jurassic AC,
separated by a subhorizontal fault. The size of this
klippe is 5 km east-west, 3 km north-south and
0.5 km in thickness. In addition, there are several
supplementary examples to support the klippe-
style occurrence of the Kurosegawa rocks else-
where in Southwest Japan (Suzuki et al. 1990;
Isozaki et al. 1992; Suzuki & Itaya 1994).
Occupying the same structural horizon and shar-
ing common geologic history, the pre-Jurassic
rocks of the Inner Zone and the Kurosegawa Belt
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12 Y. Isoxaki
Table 1 Comparison of components between the Kurosegawa Belt of the Outer Zone and pre-Jurassic belts in the Inner Zone,
Southwest Japan (Isozaki & Itaya 1991)
Compared items
Rocks and strata
400-Ma granltes
Siluro-Devonian
Ultramafic rocks
Paleozoic calc-alkali volcanic rocks
Schists
280-350 Ma
170-230 Ma
End-Permian AC
Faunal endemism
L. kumaeiisis As. (fusulinid)
l? bipartitus-F. charveti As.
(radiolaria)
Tectonic events
PermiadTriassic unconformity
(Akiyoshi orogeny) Pre-Jurassics of the Inner Zone
Dai orthogneiess (Ng)
Yoshiki and Fukuji F. (Hm)
Oeyama and Yakuno ophiolties (Mz)
Higashihirano F. (Ng)
Sorayama F. (Hm)
Nishiki G.
Ise schists (Hm)
Renge schists
‘Sangun’ schists
Nishiki G. (Ak)
Perman formations in the Akiyoshi
and Taishaku areas
Maizuru G. (Mz)
Kozuki and Oi F. (Ut)
Mine unconformity (Tsunemori/Mine) Kurosegawa components
~~~~
Mitaki granites
Okanaro G.
Serpentinite at Ino
Gokasho-Arashima T. L.
Okanaro G.
Shingal F.
Ino F.
Sawadani schists
Ino F. Agekura F.,
Agawa Unit,
Kamikatsu phyllites
Shingai F.
‘Shirakidani G.’
Sawadani G.
Kuma F.
Shingai F.
Kuma F., Shingai F.
Takano F.
Sakashu unconformity
(Hisone/Kochigatani)
AC, accretionary complex; Ng, Nagato tectonic zone along the Oki belt; Hm, marginal part of the Hida belt. I I
Fig. 10 Example of the Kurosegawa klippe in
the Nakatsu area, Southwest Japan (Isozaki &
ltaya 1991) A component of the Kurosegawa
belt (180-230-Ma semi-schists unit) occurs
as a klippe tectonically resting on the Jurassic
accretionary complex (AC) unit (nappe) of the
Chichibu belt
are regarded as having once formed different parts unit. The erosion level of the nappe of the pre-
of the same nappe, although they are at present Jurassic rocks is not deep enough in northern
isolated by 100-150 km in a north-south direction Kyushu to expose the underlying Jurassic nappe
in central Japan. In this model, the Jurassic AC in nor to form an isolated klippe/window (Fig. 11) so
the Mino-Tanba belt and Chichibu belt, on both that we can observe the merging of the two belts
sides of the Kurosegawa belt (Fig. 3a), are accord- into one with the greatest width of the Pre-Jurassic
ingly regarded as a structurally underlying nappe rocks in Kyushu (Fig. 2). The long-term exposure
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Accretion and collision in Permo-Triassic Japan 13
essary, because the Zone-bounding Median Tec-
tonic Line (MTL) initially became activated after
the Late Cretaceous (Isozaki 1996). urosegawa belt
~yu~h'
Sh,kOku
Fig. 11 Model showing 3-D configuration of the Kurosegawa klippe and the
relation between the erosion levi:l and the surface trajectory of the base of the
klippe (Isozaki et a/ 1992) Compare the widest distribution area of the
Kurosegawa belt in Kyushu by shallow erosion and its absence in !he central
Kii Peninsula by deep erosion
of the pre-Jurassic nappe at the surface (occupying
the top horizon of piled nappe) is supported by the
occurrence of non-metamorphosed Mesozoic
shallow-water clastics that unconformably cover
the pre-Jurassic rocks both in the Inner Zone and
the Kurosegawa belt. The chaotically mixed nature
of the Kurosegawa. belt may have developed
through the gravitational collapse of the nappe
along an advancing nappe front.
The biggest advantage of the current klippe
model lies in the following aspects. This model
allows us to reconstruct paleogeography with only
one subduction zone (probably along the Yangtze
continental margin), while previous models, in par-
ticular collision models, require more than two
subduction zones (i.e. at least one orogenic com-
plex for the Inner Zone and another identical
orogenic complex for the Outer Zone at the same
time). It is unlikely to assume that two or more
subduction systems generated identical AC and
high-PIT metamorphosed AC simultaneously in the
same region. Furthermore, there is no reason to
assume a hypothetical island arc nor microconti-
nent, because sialic rocks in the Kurosegawa belt
are present in quite minor amounts. Judging from
the OPS analysis, the Permian AC and 220-Ma
high-PIT schists are products of the subduction of
a major oceanic plate of -100 million years old
(probably the Farallon Plate). Short-lived marginal
basins cannot produce such an OPS. The recogni-
tion of the Kurosegawa klippe has brought about a
new model of subhorizontally layered nappes for
the whole Late Paleozoic-Mesozoic orogenic his-
tory of Japan. Accordingly, the distinction of the
Inner and Outer Zones of Southwest Japan in
pre-Cretaceous tectonics was revealed to be unnec- PERMO-TRIASSIC COLLISION COMPLEXES IN JAPAN
In contrast to the abovementioned AC units, devel-
opment of the medium-pressure-type gneidschist
units represents a quite distinct facet in the tec-
tonic history of Permo-Triassic proto-Japan. These
units are regarded as a product of the collision
between the Yangtze and Sino-Korean cratons,
instead of the AC. The metamorphic grades of
these units reached the amphibolite to granulite
facies, as evidenced by kyanite, staurolite, and
sillimanite. Such a metamorphic aspect is rare in
other units. The protoliths include terrigenous clas-
tic sediments, mafic to acidic volcanic rocks of an
alkaline nature, and bedded impure carbonates,
lacking bedded chert. Such a lithologic assemblage
is unique and remarkably different to those of the
AC units in Japan. In order to emphasize the
contrast with the AC units, I briefly describe the
Permo-Triassic orogenic units of the Hida and Oki
Belts in Southwest Japan.
HlDA BELT
Facing the Japan Sea to the north, the Hida belt
(Fig. 2) occurs in the most continent-sided position
in Japan. The rocks of the Hida belt occupy as a
nappe the highest structural horizon in Southwest
Japan (Komatsu 1990; Fig. 12), which thrust ob-
liquely upon the Permian and Jurassic AC around
the end of the Jurassic (Sohma & Kunugiza 1993).
The main components of the Hida belt are high-
grade gneisses associated with granitic rocks. In
addition, fragments of Paleozoic shelf sediments,
Paleozoic high-PIT schists, Permian AC and ser-
pentinite occur along its eastern and southern
margin. Polymetamorphism is detected in terms of
petrography and a trident clustering in radiometric
ages at 330, 240 and 170 Ma (Sohma et al. 1990;
Suzuki & Adachi 1991). The 240-Ma event corre-
sponds to the timing of amphibolite facies meta-
morphism and accompanied migmatite formation.
In addition to the regional development of a
medium-pressure-type schist belt (Hiroi 1983), lo-
calized occurrences of kyanite, staurolite and silli-
manite grains were recently reported from river
sands exclusively in the central gneiss domain
(Kano et al. 1993). This medium-pressure-type
metamorphism is regarded as a consequence of a
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14 Y. Isoxaki
230 Ma medium pressure-type metamorphic rocks a 200-170 Ma granile
Fig. 12 Geologic framework of the Hida belt in Southwest Japan. (a)
Geotectonic sketch map of central Japan showing the overlapping relationship
of the Hida belt (nappe) to the other belts of Southwest Japan (modified from
Komatsu 1990). Ak, Akiyoshi belt; Sn, Sangun belt; Mz, Maizuru belt; MT,
Mino-Tanba belt; Ry, Ryoke belt. Note a linear magnetic anomaly zone
(hatched area) beneath the Hida belt suggesting an easterly extension of the
Maizuru belt. (b) Geologic map of the Hida belt (modified from Sohma 8,
Kunugiza 1993) Note the marginal thrust zone composed of fragmentary
bodies of geotectonic units in Southwest Japan including fragments of the
Oeyama belt (Oe), Renge belt (Rn), Akiyoshi belt (Ak), and Maizuru belt (Mz)
continent-continent collision between the Sino-
Korean and Yangtze blocks (Sohma et al. 1990;
Sohma & Kunugiza 1993).
The gneiss protoliths contain thick impure car-
bonates intercalated with mafic to acidic volcanic
rocks of an alkali rock series and subordinate
peraluminous argillaceous sedimentary rocks yield-
ing chloritoid (Sohma et al. 1990). Late Carbonif-
erous fossils (foraminifera and bryozoa) were re-
ported from less-metamorphosed marble (Hiroi et
al. 1978). The absence of chert, the occurrence of
peraluminous pelitic rocks, and the coexistence of
bimodal volcanic rocks suggest that the protoliths
were likely to be derived from the Middle-Late Paleozoic continental shelf-platform sediments ac-
cumulated on a rifted continental margin (Sohma
et aL. 1990). The Permian brachiopod fauna from
the marginal part of the Hida belt contain mostly
cold-water (Boreal) elements (Nakamura & Tazawa
1990), probably indicating its northerly position
with respect to other Tethyan fauna from the
low-latitude mid-Tethyan realm.
The Hitachi-Takanuki belt (Fig. 2) is the only
known area of kyanite-bearing 250-Ma metamor-
phic rocks in Northeast Japan (Tagiri 1973; Tagiri
et aL. 1992). Its protolith assemblage, fossil ages
and U-Pb zircon age (200-280 Ma; Hiroi et al.
1994) all suggest that this unit probably represents
not only an eastern extension of the Hida Belt in
Japan but also a fragment of the 250-Ma collision
complex in central China and Korea.
OK1 BELT
Isolated physiographically and geologically from
the Hida Belt, the Oki Belt exposes another
non-AC unit in Southwest Japan (Fig. 2). The
distribution of component rocks is limited to the
Oki Islands in the southern Japan Sea where a
high-grade gneiss complex occurs (Hoshino 1979).
The complex consists mostly of psammitic and
pelitic gneisses with minor amounts of amphibolite
and marble. Regarding protolith composition, this
lithologic assemblage differs considerably from that
of the Hida Belt. Metamorphism reached the am-
phibolite facies and went partly into the granulite
facies. The age of this metamorphic event is dated
at 250 Ma (Suzuki & Adachi 1994; Fig. 13). In
addition, inherited zircon grains with ages of 3.0,
2.0, 1.7, and 1.25 Ga and 350 Ma have been
identified, indicating that protoliths formed adja-
cent to an Archean-Proterozoic continental block.
Like the Hida gneiss, the lithologic assemblage
of the Oki gneiss suggests the protoliths were
derived from a rifted continental margin, but the
latter lacks an extensive carbonate platform. Al-
though the Hida and Oki Belts share the same
metamorphic age of 250 Ma, differences in proto-
lith lithology and faunal provinciality suggest that
these two belts may have been derived from
different continental margins, and that these two
margins were primarily separated from each other.
The Southern Kitakami belt in Northeast Japan
(Fig. 2) is regarded as an eastern extension of the
Yangtze block, on the basis of stratigraphy of
Middle-Late Paleozoic continental shelf facies sed-
iments with an affinity to that of the Yangtze
platform rather than Sino-Korea (Isozaki &
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Accretion and collision in Permo-Triassic Japan 15
0.2-
0.1 -
-
-0
"'01 M01 from sample 1404A M71 from sample 0304
250+_20 Ma 250-120 Ma
MSWD=O.IO MSWD=0.26
1, I I , , I
Maruyama 1991). The previously accepted correla-
tion between the Southern Kitakami belt and the
Sino-Korean block is refuted by the total absence
of Siluro-Devonian strata in the latter. The faunal
analysis on Permian bivalves positively supports
this correlation (Nakazawa 1991), on the other
hand, those of Permian corals or brachiopods (Kato
1990; Nakamura & Tazawa 1990) appear equivo-
cal. Judging from geological and paleontological
similarities, the Oki belt is correlated not only to
the Yangtze block but also to the Southern
Kitakami belt (Isozalii & Maruyama 1991). e 0.2- M30 from sample 1404C 0.2-
25OSO Ma
MSWD=0.44
0 1- 0.1- Fig. 13 Th-U-total Pb isochron of mona-
zite grains from the Oki gneiss (Suzuki &
Adachi 1994) (a-c) lsochrons for a single
monazite grain; (d) isochron for 35 monazite
grains from the same rock sample. Note the 01 I I I I I I I Ot
250-MA GRANITE 35 grains from sample 1502
250i10 Ma
MSWD=0.54
I, I I, I I I I
There are several scattered occurrences of 250-Ma
granites in Japan, although minor in volume: they
are usually less than a kilometer in diameter.
Besides the Maizuru granite contained in the
Yakuno ophiolite, most of them occur along the
MTL on its Outer Zone side, forming a nappel
klippe upon the younger Jurassic-Cretaceous AC
and/or their metarnorphosed equivalents (On0
1983; Takagi & Fujiimori 1989). Their field occur-
rence suggests subhorizontal transport through
nappe tectonics from the Inner Zone, however, the
root zone of these granite nappes has not been
identified. From the viewpoint of regional tectonics,
there are two alternatives for the origin of the
250-Ma granite: one related to subduction as being
responsible for forming the Akiyoshi-Maizuru AC and Sangun meta-AC, and the other related to the
collision between the Yangtze and Sino-Korean
cratons. No critical evidence has yet been obtained,
however, the latter option appears likely at present
because 250-Ma granites are found in the Hida and
Southern Kitakami belts (Suzuki et al. 1992), in
particular, associated with high-grade gneisses in
the Hida belt, while no large granite batholith of ca
250 Ma has been found in Japan.
DISCUSSION
On the basis of the above descriptions, this section
summarizes significant tectonic aspects of Permo-
Triassic Japan. First, tectonic implications of
subduction-accretion of oceanic topographic highs
at trenches and the subhorizontal nature of
accretion-generated orogens are discussed. Then
the unique tectonic setting featuring two 250-Ma
orogens of distinct characteristics are discussed
with respect to the paleogeographic reconstruction
and faunal provinciality of western Panthalassa at
that time.
ACCRETING EFFICIENCY OF SEAMOUNTS AND
OCEANIC PLATEAUS
Compared with the Jurassic and Cretaceous AC in
Japan (Isozaki 1997; Kimura 1997), the Permian
AC in Southwest Japan are unique in having
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16 Y. Isoxaki
abundant oceanic materials; that is, igneous and
sedimentary rocks derived primarily from oceanic
crust such as basaltic/gabbroic greenstones, reef
limestone, ophiolitic complexes including ultramafic
rocks, and deep-sea chert. In particular, the pre-
dominance of large slabs/blocks of reef limestone
and ophiolitic rocks is remarkable when compared
with other ancient AC exposed in Japan. Despite
the considerable variation in mutual volume ratio,
oceanic materials usually do not exceed 30% vol-
ume of all components of an AC (Isozaki et al.
1990). The Permian AC in Southwest Japan as a
whole also follow this empirical rule but the local-
ized occurrence of considerably huge oceanic blocks
appears exceptional; for example, the Akiyoshi reef
limestone is 10 x 10 km wide and 1 km thick, and
the Yakuno ophiolite is 100 km long and 8 km thick.
Lithostratigraphic and petro-chemical analyses
suggest that such blocks are allochthonous frag-
ments of paleoseamounts and paleo-oceanic pla-
teaus. Given the average bathymetry of the mod-
ern ocean, the height of the reef-capped Akiyoshi
paleoseamount is estimated at 4 km, even allowing
for reef-drowning after the over 700 m-thick lime-
stone accumulation under photosynthetic shallow
water (Fig. 14). The scarcity of basaltic green-
Trench A”;” ,!>.,- -- stones underlying reef limestone indicates that a
basaltic seamount pedestal more than 3 km thick
per se has vanished to the subduction zone without
accretion, just leaving its reef cap in the trench.
Concerning the Maizuru paleo-oceanic plateau,
the original thickness of total crustal materials is
estimated to be at least 15 km. On the other hand,
the on-land exposed Yakuno ophiolite is -8 km
thick, indicating that it is highly dismembered or
telescoped. This mismatch in thickness suggests
that more than 7 km of crustal materials have been
separated from the original oceanic crust. These
examples of the Permian AC support the sugges-
tion that large oceanic reliefs (seamount, rise, or
plateau) can subduct at a trench, leaving thin
fragments in AC that are thinner and smaller than
the primary thickness and volume (Isozaki et al.
1990). What appears contradictory is the scarcity
of accretion of basaltic greenstones from a highly
elevated seamount while there is abundant accre-
tion from a rather low-lying oceanic plateau. This
contrast may have been due to the difference in
total volume of topographic relief that entered the
ancient trench, or to differences in rigidity because
the oceanic plateau is larger in volume than iso-
lated seamounts in general (Fig. 14).
Mid-Oceanic Ridge
201 II II
Fig. 14 Scheniatic diagram showing a crustal section of an ancient oceanic plate reconstructed on the basis of the stratigraphy and metamorphic constraints from
the Maizuru ophiolite and Akiyoshi limestone within the Permian accretionary complex (AC) in Southwest Japan Note that this reconstruction is highly speculative
and needs further testing with information from modern examples when it is available Prominent oceanic topographic-highs show the Akiyoshi paleoseamount with
an over-700 m-thick reef carbonates above the carbonate compensation depth (CCD) and the Maizuru paleo-oceanic plateau with an over-1 5 km-thick crust below
the CCD
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Accretion and collision in Permo-Triassic Japan 17
On the other hand, collision-subduction of a
seamount is regarded as the main mechanism of
subduction erosion by the destruction of previously
accreted materials in the trench inner wall (von
Huene & Scholle 1991). Accordingly, the total
volume of the originally formed Permian AC was
probably much greater than what we see today,
and a considerable amount of AC material was
subducted along with collided seamounts. The cha-
otically mixed units in the Permian AC probably
represent re-accreted AC material after the pri-
mary accretion and collision-induced collapse, and
their predominance a.lso supports the tectonic ero-
sion event.
The abovementioned abundance of material de-
rived from paleoseainounts and plateaus implies
that the subducting oceanic plate may have carried
numerous topographic highs on it. Furthermore, it
indirectly suggests that an ancient plume (or su-
perplume) somewhere in the Carboniferous-Early
Permian Pacific/Panl,halassa Ocean produced nu-
merous topographic highs, like the one in the
Cretaceous mid-Pacific. This can be tested in coeval
AC along the eastern Panthalassic margin by
future work.
SUBHORIZONTAL STRUClURE OF ACCRETIONARY OROGEN
Another important contribution from the latest
studies on the Permo-Triassic AC orogen in South-
west Japan is the documenting of a predominant
piled nappe structure, including isolated klippes
and windows. The Jurassic AC also contributed to
the piled nappe structure, and these suggest that
the subhorizontal structure governs the entire
Paleozoic-Mesozoic orogens of Southwest and
Northeast Japan and Ryukyus (Isozaki 1996,
1997). With respect to the piled nappe structure,
the downward-younging growth polarity is ob-
served particularly in Southwest Japan (Fig. 3).
The uni-directional-younging polarity implies that
ancient AC have accumulated downward, keeping
essentially the same growth direction in accordance
with the initial accretion polarity at the trench on a
smaller scale. Occurrence of similar AC have been
reported sporadically from various areas in the
world but there has been no better documented
example of such a subhorizontal structure in an AC
orogen as in Japan, which stretches more than
2000 km in length and 200 km in width.
Concerning the emplacement mechanism of each
AC nappe into the piled nappe edifice, active move-
ment of the lowermost nappe is emphasized here.
With respect to the continent, it is not an overlying nappe but an underlying nappe that moves horizon-
tally to make piled nappes because nappe-forming
layer-parallel shortening is driven by oceanic sub-
duction. This suggests that the pre-Jurassic com-
plex stayed mostly in the same position relative to
the continent, and without long-distance transpor-
tation, they turned into an ‘autochthonous’ nappe
when the younger Jurassic AC nappe was em-
placed underneath it. In other words, the Kuroseg-
awa klippe (or nappe) per se did not travel for
more than 100 km across-arc, all the way from the
present Inner Zone of Southwest Japan.
Subhorizontal piled nappe structures have been
previously emphasized for continent-continent
collision-related orogens such as the Himalayas
and Alps, as well as foreland fold and thrust belt
such as the Canadian Rockies. On the other hand,
concerning the subduction-related orogens classi-
fied as Cordilleran-type by Dewey & Bird (1970),
vertical structures have been emphasized in the
central or aged part of an orogen. In the 1980s,
the ‘exotic terrane’ concept also over-emphasized
vertical faults which bounded terranes by second-
ary strike-slip movement. Analyses of other AC-
dominant orogens will test whether or not a sub-
horizontal nature is essential to oceanic sub-
duction-related orogens in general.
The intrusion of 85-Ma I-type granitoids into the
220-Ma high-P/T Sangun nappe provides a refer-
ence for estimating the rate of oceanward growth
of the subhorizontal AC-dominant orogen in Japan.
The granite penetrating the 220-Ma schists repre-
sents a part of a large Late Cretaceous granitic
belt called the Ryoke belt in Southwest Japan,
which is regarded as the basement of a Late
Cretaceous volcanic arc. In general, the volcanic
arc occurs -100-200 km inland from the coeval
trench. As the oceanward margin of the 220-Ma
Sangun nappe approximately represents a paleop-
osition of the 220-Ma trench, the oceanward shift
of the trench position within the AC-dominant
orogen is suggested to be 100-200 km during an
approximately 135-million-year (220-85 Ma) inter-
val, which appears consistent with the average
horizontal growth rate of the Japan orogen, that is,
400 km in 400 million years (Isozaki 1996).
COLLISION SUTURE AND LATERAL VARIATION
The occurrence of an ultrahigh-pressure metamor-
phic belt in East Asia is restricted to the Dabie-
Sulu zone (Cong & Wang 1995), particularly to
the central bending part of this collision suture
between the Yangtze and Sino-Korean blocks
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18 Y. Isoxaki
(Fig. 15). The metamorphic grade of the collision-
related units apparently decreases westward, and
probably eastward judging from the absence of an
ultrahigh-pressure unit to the east of the Shangdon
Peninsula, Northeast China. The 250-Ma medium-
pressure-type metamorphic units in the Oki and
Hida belts in Southwest Japan are the best candi-
date for the easterly extension of the suture, as the
rest of Southwest Japan is occupied by non-
collisional orogenic units. The Hitachi-Takanuki
belt and Southern Kitakami belt in Northeast
Japan can be viewed as the probable easternmost
tract of the suture. Restoring their strike-slip
displacement induced by the Miocene Japan Sea
opening (Otofuji 1996), the pre-Miocene paleoposi-
tion of these fragments is expected to be at a
higher latitude than their present position.
The decrease in metamorphic grade toward both
directions along the suture may demonstrate that a
considerable lateral variation on a scale of more
than 1000 km existed within the same collisional
orogen. Major compressional stress by collision
probably had been localized in the central portion
;ion orogen
subduction zone
terrigenous cIastics
/
Fig. 15 Triassic-Jurassic Japan The Qinling-Dabie zone in China repre-
sents the collision suture between the Sino-Korean (North China) and Yangtze
(South China) continental blocks (modified from lsozaki & Maruyama 1991 )
Note the concentrated occurrence of coesite/diamond-bearing ultrahigh-
pressure metamorphic rocks (shown by open squares after Cong & Wang
1995) exclusively in the bending part of the suture in the Dabie-Shandon
areas See text for information regarding autocannibalism of the Yangtze Block
related to the unique double-edged orogenic setting enough to bring buoyant continental materials
down to -100 km deep, while other parts suffered
less stress and metamorphism. This localization of
maximum stress may have been driven by the
collision of two continental margins with heteroge-
neous geometry. Continental collision usually starts
from the promontory-promontory collision stage as
perceived earlier by Dewey & Burke (1974). Ac-
cordingly, the ultrahigh-pressure part within a
collision suture may indicate the ancient site of
promontory-promontory (or promontory on one
side) collision (Fig. 16). Concerning the detailed
exhumation history of this ultrahigh-pressure unit,
Maruyama et al. (1994) proposed a model includ-
ing subduction of microcontinent and subsequent
delamination of a subducted slab. Their microcon-
tinent can be virtually replaced by the detached
continental fragment of a collided promontory.
However, lesser metamorphic grades in lateral
extensions probably suggest that the compres-
sional stress was at a maximum around the inden-
tor (promontory), while less stress was experienced
on both sides of the indentor, at least not enough
stress to bring buoyant protoliths down to such
deep levels. Concerning this along-suture metamor-
phic variation, another explanation is possible;
namely, that collision-induced escape tectonics may
have occurred to extrude excess materials laterally
to free spaces of oceanic domain and to disrupt
primary metamorphic zoning (Fig. 16b), just like
the case of the Indochina block induced by the
Indian collision against Asia (Tapponnier et al.
1982). The 230-Ma right-lateral shearing detected
in the Korean Peninsula (Yanai et al. 1985; Cluzel
et al. 1991) and in the Hida belt (Nagase et al.
1990; Takagi & Hara 1994) may represent a part
of a conjugate fault system related to escape
tectonics in the eastern marginal part of the colli-
sion suture. The present nappe occurrence of the
Hida Belt (Komatsu 1990; Fig. 12) and its forma-
tion process is not discussed here because it is a
result of the endJurassic across-arc contraction
(Sohma & Kunugiza 1993) which occurred long
after the 250-Ma collision event.
The correlation between the two cratons in
China and the geotectonic units in Japan presented
here is in remarkable contrast to previous theories.
The HidaISouthern Kitakami belts have been com-
monly correlated to the Sino-Korean block simply
because they show a physiological proximity at
present (Kobayashi 1941; Minato et al. 1965;
Taira & Tashiro 1987). The latest tectonic evalua-
tion of the suture between the twa-cratons and the
search for its easterly extension in Japan, however,
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Accretion and collision in Permo-Triassic Japan 19
tinct types side by side; the Akiyoshi-Sangun-
Maizuru accretionary orogen and the Hida-Oki
collisional orogen, and both had extensive dimen-
sions with lengths of some 1000 km. From the
viewpoint of the regional tectonics of Asia, what
appears more important is the collision-related
orogenic belt between the Yangtze and Sino-
Korean continental blocks because these two form
the core of present East Asia, which grew much
bigger through the collision/amalgamation with
Siberia, Bureya (Amuria) and other continental
pieces originally detached from Rodinia at 750-
700 Ma.
The AC-related orogen facing the proto-Pacific,
however, has nearly the same weight in geotec-
tonic significance as well, because this involved the
subduction of a major oceanic plate in the proto-
Pacific, probably the Farallon Plate of some thou-
sands of kilometers wide. This implies that the
subduction of a major plate continued actively
along the western margin of Panthalassa, at least
along the Yangtze margin, even at the time of the
Pangea assembly when most of the subduction
zones between continents were terminated by col-
lision. In short, the Yangtze craton was double-
edged in the Permo-Triassic, featuring two active
fronts of completely different natures at the same
time (Fig. 16b).
Such a double-edged situation brought another
unique geologic setting around Japan in the follow-
ing Jurassic time. The collision-related exhumation
of ultrahigh-pressure rocks probably triggered a
regional upheaval along the suture and it resulted
in the provision of abundant terrigenous clastics
along the suture, away from the center to the
marginal delta. The collision-derived clastic mate-
rials are delivered eastward along the suture to
build a delta (Fig. 15). By transport further along
the trench axis, they consequently returned to the
opposite side (southeastern margin) of the Yangtze
Block and again accreted to the continent (Isozaki
& Maruyama 1991; Isozaki 1997). Thus semi-
autocannibalistic-material recycling occurred
around the Yangtze Block.
The paleogeographic reconstruction of Permo-
Triassic proto-Japan and its vicinity (Fig. 16) can
be tested by the faunal provinciality of Permo-
Triassic taxa such as bivalves, fusulinids and bra-
chiopods. In general, distinction of faunal provinces
is apparent during the Permian (Ross & Ross 1983;
Ishii 1990), while it is ambiguous in the Triassic
due to its monotonous nature. According to Naka-
zawa (1991), Permian faunas from clastic cover
sequences in the Maizuru and Kurosegawa Belts Dabie Promontory
Q S Kitakami
All
1 Akiyoshi A
paleo-seamounts - Panrhalassa
Farallon Plate
Maizuru 0 paieo-oceanic plateau
.%... ,?$$
Farallon Plate
Izanagf Plate
Fig. 16 Paleogeographic reconstruction of proto-Japan and its vicinity (a)
Middle Permian ca 270 Ma klote the paleopositions of the Hida belt on the
Sino-Korean margin, Oki and Southern Kitakami belts on the Yangtze margin
and those of the Akiyoshi paleoseamount and Maizuru paleo-oceanic plateau
in Panthalassa (b) Early-Middle Triassic, 250-230 Ma Note the double-
edged orogenic nature of the Triassic Yangtze craton featuring a continental
collision against the Sino-Korean block on the north and subduction by the
proto-Pacific (Farallon) oceanic plate on the south The sinuous part of the
suture in the Dabie-Sulu area I:, characterized by an ultrahigh-pressure meta-
morphic belt probably reflectinij the site of a promontory collision between the
two continental blocks Dextral strike-slip deformation in the Hida belt and in
the Korean Peninsula suggests eastward escape tectonics driven by the
collision
suggested a strong link between the Yangtze block
and most of the Japanese units.
JUXTAPOSITION OF CONTRASTING OROGENS AND
PALEOGEOGRAPHY
In the Permo-Triassic of proto-Japan and its sur-
roundings, there were two active orogens of dis-
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20 Y. Isoxaki
show a strong affinity with the warm-water
Yangtze fauna rather than those from the Sino-
Korean Block, while the Hida and Southern Kita-
kami faunas partly accompany Boreal representa-
tives (Nakamura & Tazawa 1990; Tazawa 1992).
These observations, including the monotonous as-
pect in Triassic fauna, appear concordant with the
present paleogeographic reconstruction. The incor-
poration of some Boreal elements in the Hida and
Southern Kitakami fauna may indicate a stronger
influence of cold currents along the northwestern
Panthalassa according to their relatively northerly
paleoposition to the rest of the Southwest Japan
fauna. The Hida, Oki, and Southern Kitakami belts
have previously been correlated with the Sino-
Korean block (Kobayashi 1941; Minato et al. 1965;
Tazawa 1992), however, this correlation faces dif-
ficulty in identifying the suture between the Sino-
Korea and Yangtze blocks in and around Japan
unless an arc-parallel strike-slip displacement (of
an unrealistically large scale) is assumed.
On the other hand, the Akiyoshi reef limestone is
characterized by a peculiar faunal succession, dis-
tinct from those from the Hida and Southern
Kitakami belt. The Early Carboniferous fauna from
the Akiyoshi limestone is similar to Australian
fauna in the southern hemisphere (Kato 1990). The
OPS analysis on the Akiyoshi AC suggests that the
Carboniferous paleoposition of the Akiyoshi paleo-
seamount with reef was some thousands of kilome-
ters away from the Yangtze coast, that is, some-
where in the open ocean within western
Panthalassa. The Tethyan aspects of the Akiyoshi
limestone, particularly in the Permian part, were
secondarily overprinted when the paleoseamount
entered the paleo-equatorial domain immediately
before the accretion (Fig. 16b). Thus its peculiar
faunal succession and distinction from those in the
Hida and Southern Kitakami Belts is consistent
with the paleogeographic reconstruction based on
accretion tectonics.
CONCLUSION
The latest views on the Permo-Triassic tectonics of
the Japanese Islands are summarized as follows.
(1) The Japanese Islands record two coeval and
contrasting types of convergent orogeny in the
Permo-Triassic time: an accretionary orogeny be-
tween the Yangtze craton and subducted paleo-
Pacific seafloor (probably the Farallon Plate) and a
collisional orogeny between the Yangtze and Sino-
Korean cratons. (2) The Permo-Triassic accretionary orogeny in
Japan involved collision-subduction of a paleosea-
mount and paleo-oceanic plateau, and left smaller/
thinner fragments in the accretionary complex.
(3) A subhorizontal piled nappe structure gov-
erns the Permo-Triassic orogen in Japan, showing
several klippes and a peculiar ‘sandwich structure’
of high-P/T schist nappe.
(4) The Qinling-Dabie suture in central China, a
major collisional suture in Asia between the
Yangtze and Sino-Korean cratons, extends east-
ward to the Hida and Oki belts in Southwest Japan
and the Hitachi-Takanuki belts in Northeast
Japan.
(5) A paleogeographic reconstruction around
Permo-Triassic proto-Japan with the double-edged
Yangtze Block is consistent with the faunal char-
acteristics of Permo-Triassic units in Japan.
ACKNOWLEDGEMENTS
The present author thanks S. Maruyama for his
comment on ophiolite, and W. McDonough, A.S.
Jayko and A. Taira who read the manuscript and
gave valuable comments.
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Isozaki 1997 Island Arc - March 1997 - Isozaki - Contrasting two types of orogen in Permo_Triassic Japan Accretionary versus (1).txt
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A u t h o r ( s ) Y A M A M O T O , A k i h i k o
C i t a t i o n J o u r n a l o f t h e F a c u l t y o f S c i e n c e , H o k k a i d o U n i v e r s i t y . S e r i e s 7 , G e o p h y s i c s , 1 2 ( 2 ) , 7 5 - 9 5
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Jour. Fac. Sci., Hokkaido Univ., Ser. VII (Geophysics), Vol. 12, No.2, 75-95, 2004.
Dense Clustering of Latest Cenozoic Caldera-like
Basins of Central Hokkaido, Japan,
Evidenced by Gravimetric Study
Akihiko Yamamoto
Institute of Seismology and Volcanology, Graduate School of Science,
Hokkaido University, Sapporo, Japan
(Received January 9, 2004) )
Abstract
This paper examines gravity structures of the central Hokkaido using new
gravity data, and attempts to present preliminary interpretations of their characteris
tic features, particularly with special attention to the numerous sedimentary basins of
latest Cenozoic age (late Miocene to Quaternary). We compile more than 6,000
gravity data of the study area and produce a new gravity anomaly map. The new
map in the central Hokkaido delineates two characteristic features that are divided
by a boundary which falls approximately along the Tokoro Tectonic Line (TTL) of
the map area. To the east of TTL, the Bouguer anomaly field is characterized by the
high-amplitude gravity ridge attributable to Mesozoic sequences associated with
several ellipsoidal gravity depressions having an almost NE-SW major axis of 15~20
km, whereas to the west, anomaly relief is much lower with several closed gravity
depressions relative to the region to the east of TTL. It is quite intriguing that these
remarkable lows, dominant over the mountainous area to the west of TTL, clearly
form a dense cluster of closed depression with a diameter of about 10 km, which well
correlates with the distributions of the known basins. These features are much
strengthened by a relief-shaded Bouguer image. Gravity analysis by horizontal
derivatives and high-pass filtering of the gravity field shows that the major high
Bouguer gradient zones, found to be nearly closed in a ring-shaped or oval-shaped
form, are remarkable in the central part of the map area. Several of them indicate
a better coincidence between the basin rims and the location of a steep gravity change.
This implies that they have caldera-like collapse structure, and that the subsurface
part of these basins accords with a steep-sided depression, with a flat bottom,
probably filled with low-density volcanic materials. The Bouguer anomaly contours
around these caldera-like basins are likely those of the thickness contours of the basin
fill, which yields significant constraints on the subsurface structure beneath the basin.
Assuming the average density of the basin fill to be 2.2g/cm3 (a density contrast of
0.47 g/ cm3) in one of the typical Cenozoic basins in the central Hokkaido, we obtain
the result that an apparent maximum thickness of the basin fill amounts to be 750 m
~1,000 m, which corresponds to a gravity attraction with a relative amplitude of 15
and 20 mgal, respectively.
76 A. Yamamoto
1. Introduction
The central Hokkaido, Japan, is characterized by large mountainous areas
with notable calderas and volcanic plateau underlain by tuffs and lavas, as well
as numerous small-scale sedimentary basins (e.g., Yamagishi, 1976; Oka,
1986; Kato et al., 1990) which have continuously undergone tectonic move
ments of east-west compression (Miyasaka and Matsui, 1986). Recent studies
show that these sedimentary basins have been formed during late Cenozoic (late
Miocene to Quaternary) age and their tectonic evolution is roughly explained in
terms of plate tectonics (Oka, 1986). Gravimetric studies are becoming
widely acknowledged as a powerful and useful tool for studying the distribu
tions of subsurface masses that may be associated with tectonic or volcanic
activity to form such basins or calderas. The gravity field prevailing over the
central Hokkaido, particularly around the Cenozoic basins, is best reflected on
Bouguer gravity anomalies that are most sensitive to the subsurface mass
distributions and are influenced by the nature of underlying geological forma
tions.
A number of gravity studies have given significant constraints on basins or
calderas around volcanic areas. Healey (1968) described the gravity anomaly
at Pahute Mesa, located north of Timber Mountain, and attributed the anomaly
to the volcanic rocks of the Silent Canyon caldera complex. The gravity
anomaly he described is roughly circular and its overall appearance has prob
ably affected the thinking of many regarding the geometry of the caldera
complex. Later Ferguson et al. (1994), based on gravity and seismic travel
time data, showed that the Silent Canyon caldera complex is actually a nested
set of buried calderas, and that abrupt changes in the subsurface thickness of the
caldera-forming units occur across the faults, indicating that these linear fea
tures served as caldera boundaries. Gettings and Griscom (1988) draw a
conclusion that a ring dyke of mafic composition is inferred to intrude to near
surface levels along the caldera ring fractures, and low-density fill of the caldera
floor probably has a thickness of 0.7-0.9 km from gravity modeling study of
Newberry Volcano in U.S.A. Eaton et al. (1975) and Kane et al. (1976) gave
interpretations of gravity anomalies over a young volcanic center in connection
with Yellowstone and Long Valley Caldera in U.S.A., respectively. The pres
ence of five major centers in the Taupo Volcanic Zone (New Zealand), each
associated with an approximately circular negative anomaly, was revealed by
gravity analysis (Rogan, 1982). Carle (1988) suggested that Long Valley
Caldera has a piston-shaped collapse structure based on three-dimensional
Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 77
/ Fault
/" Q(~i~~f;~~tsu T. Z.) 0,-, __ 1..L? __ Z?km
f::::. Volcanoes
(Late Pleistocene toRc:cent)
Fig. 1. Distributions of the latest Cenozoic basins around the Kitami Mountain area,
central Hokkaido (After Ok a, 1986). The perimeters of these sixteen basins are
digitized and used in relation to gravimetric ilwestigation in this study.
gravity modeling. Iterative 3-D modeling using gravity data also placed
significant constraints on the structure and volcanic evolution of Tenerife,
Canary Islands (Ablay and Kearey, 2000). Recently, Komuro et al. (2002)
reported that the subsurface structure of the Teragi Cauldron in Sw Japan may
be square shaped, with a steep rim and a flat floor, based on the fact that the low
gravity anomaly over the cauldron shows a pan-shaped gravity depression.
Understanding tectonic basins or volcanic calderas requires passive geo
physical study, as well as detailed geological and petrological studies. The less
seismicity and the absence of ongoing eruptions around the areas of the late
Cenozoic caldera-like basins (Figs. 1 and 2) have precluded geophysical inter
pretations or investigations. Although a general view of Bouguer anomaly in
the central Hokkaido shows that (1) low anomalies corresponding to Quater
nary mafic volcanic rocks that are dominant over the mountainous areas (the
Ishikari Mountains), and (2) Mesozoic high-density rocks (late Jurassic vol-
78 A. Yamamoto
canics and late Cretaceous plutonic rocks) are characterized by high Bouguer
anomalies, detailed gravity structure particularly around the basins in the
central Hokkaido has not been studied so far.
The aim of this paper is to present preliminary interpretations of gravity
structures of north-eastern part of the central Hokkaido with special reference
to the numerous small-scale sedimentary basins formed during latest Cenozoic
age (late Miocene to Quaternary). We also attempt to constrain the geometry
of subsurface structures by integrating geophysical data (Bouguer anomaly).
2. Geological Background
Fig. 4 illustrates the generalized geology of the central Hokkaido area with
the gravity anomaly of Fig. 3 at a 5 mgal contour interval superimposed for
reference. The distributions of the latest Cenozoic sedimentary basins (Fig. 1)
of the central Hokkaido are redrawn in Fig. 2 with gravity stations, where
sixteen basins are sequentially numbered according to Fig. 4 of Oka (1986).
The map area in this study is geologically and tectonophysically divided into
three zones: the Hidaka Zone, the Tokoro Zone and the Nemuro Zone (from
left to right in Fig. 4), each of which runs almost in NWN -SES direction across
Hokkaido. It is intriguing that the sedimentary basins of our interest cluster in
the northeastern part of the Ishikari Mountains (Daisetsu and Tokachi volcanic
areas) and the southern part of the Kitami Mountains, and that their linear
dimension as a whole is about 10 km in diameter. As shown in Figs. 2 and 4,
most of the basins (#2~#4, #9~#12 and #14~#16) distribute in the Tokoro Zone
Fig. 2. Locations of gravity stations used in this study with shaded image of digital
topography. Thick and red-dotted polygons with specific sequential numbers
(#1~#16) indicate the rims of the Cenozoic basins (see Fig. I) in the central
Hokkaido by Oka (1986). Heavy colored lines demonstrate known active faults
by Nakata and Imaizumi (2002) which indicate, red: certainly exist and location
is accurately determined, magenta: certainly exist and location is not accurate,
green: possibly exist (invisible), and blue: estimated fault lying at depths.
Location of the study region is shown in the bottom right corner on a map of
Hokkaido. Large closed triangles and squares, followed by three letters, indicate
geographical locations of major summits and cities (towns), respectively. ASA:
Mt. Asahidake, ASR: Ashoro, CKS: Mt. Chitokaniushi, EGR: Engaru, HNP:
Mt. Higashi-nupukaushinupuri, ITR: Ikutahara, KFJ: Mt. Kitamifuji, KIT:
Mt. Kitoushi, KMW: Kamikawa, KTM: Kitami, MAK: Mt. Meakan, MRD:
Mt. Muridake, MSP: Maruseppu, NIP: Mt. Nipesotsu, NKN: Mt. Nishi
kumaneshiri, NKP: Mt. Niseikaushuppe, OKT: Oketo, RBT: Rikubetsu, RSB:
Rubeshibe, SRM: Saroma, TBT: Tsubetsu, TKC: Mt. Tokachidake, TKU:
Takinoue, TOM: Mt. Tomuraushi, TSO: Mt. Teshiodake, UPP: Mt. Upepesan
ke, YBT: Yubetsu.
Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 79
km
o 10 20 30
44' OO'N
43' 30'N
143' OO'E 143' 30'E 144' OO'E
80 A. Yamamoto
which largely consists of Mesozoic rocks, that is, Late Cretaceous deposits
conglomerate and Late Jurassic dismembered ophiolite sequence. In contrast,
the remainder of the basins (#1, #5~#8, and #13) belongs to the Hidaka Zone
which is characterized by Late Cretaceous deposits resting on an undated
basement and Oligocene-Miocene metamorphic rocks. There are no conspicu
ous faults present around the basins (Figs. 2 and 4), which would contribute to
the shape or structure of the basins. In the eastern part of these basins,
Mesozoic rocks are extensively dominant. East of these basins lies the N emuro
Zone whose westernmost boundary is abruptly cut by the Abashiri Tectonic
Line (A TL) which runs through Kitami (KTM) to Rikubetsu (RBT) in Fig. 4.
The Kamishiyubetsu Tectonic Line sharply bounds these basins in their north
ernmost part. Several basins (#11 ~#13) are sharply bounded on the east by the
Tokoro Tectonic Line (TTL) which runs in NWN -SES direction passing
through Yubetsu (YBT), Engaru (EGR), Ikutahara (ITR), Rubeshibe (RSB) and
Rikubetsu (RBT). Also three basins (#14~#16) are located in Tokoro belt
which largely consists of Mesozoic rocks (Late Cretaceous deposits and Late
Jurassic ophiolite sequence), whose easternmost margin is bounded by ATL
(Kimura, 1981).
Yamagishi (1976) pointed out that most of these basins show a closed
polygon with a diameter of about 10 km associated with large-scale volcanic
activity. Stratigraphic study indicates that normal sediments of the basins
contain fair amount of volcanic products that are chiefly rhyolitic pumice flow
deposits subsequent to the accumulation of the basal conglomerates, and also
abut against the basement at high angle (Yamagishi, 1976). Consequently,
Yamagishi (1976) and Moriya (1983) draw a similar conclusion that those basins
possess caldera-like collapse structure that was formed during the age from
Neogene Tertiary to Pleistocene. Recently, Ishii (2001) and Ishii et al. (2002)
argued about the characteristic features of one of these basins, the Tokachi
Mitsumata Basin (#3 in Fig. 2), and they concluded on the basis of detailed
geological study that the Tokachi-Mitsumata Basin is an oval-shaped caldera
(about 13 km by 10 km), formed by effusion of Muka pyroclastic flow in latest
Cenozoic age (487 ± 17 ka). These geological interpretations with well-estab
lished evidences suggest that other basins may also be considered as late
Cenozoic "caldera or cauldron". The evidence to define them as "caldera",
however, is not discussed here since it is beyond the scope of this paper. One
of the chief objectives of this paper is to delineate gravity features and present
reliable evidence to constrain the geometry of subsurface structures by gravity
analysis.
Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 81
3. Gravity Data and Reductions
Recent gravimetric studies such as Geographical Survey Institute (GSI)
(1985), Komazawa et al. (1987), Kona and Furuse (1989), Geological Survey of
japan (GSJ) (1992), and Komazawa et al. (1999) have presented gravity
anomaly maps of Hokkaido district or whole Japanese Islands and delineated
characteristic features of Bouguer gravity. These maps, however, do not
illustrate detailed features of late Cenozoic basins in the central Hokkaido due
to lack of gravity data and/or the nonuniformity of station coverage. To
overcome this disadvantage we compiled gravity data from many institutes such
that station coverage would be as uniform and dense as possible. Gravity data
used in this paper consists of the gravity database by GSj (2000) and land
gravity data measured by Hokkaido University. In total, about 6,100 gravity
data are collected for gravity study around the central Hokkaido. Fig. 2 shows
the distribution of gravity data used in this study. Gravity data by GSj (2000)
are reprocessed in a uniform manner used in this study. Gravity data by
Hokkaido University were obtained by a LaCoste & Romberg Model Gland
gravity meter and a Scintrex gravity meter. These data were fully corrected
for latitudinal and elevation effects, instrumental height, earth tide, and secular
drift of the spring. Characteristic features and behaviors of the above gravity
meters are described by Yamamoto et al. (200la,2001b). These gravity data
were then reduced for free-air and Bouguer corrections, as well as for terrain
effects. On a mature active, volcanic areas such as the central Hokkaido,
where the topography is extreme and elevation varies in excess of 2.3 km above
and below sea level, terrain effects are large and show considerable spatial
variations. Accordingly, in this study, we compute terrain corrections using
digital elevation model spaced at every 50 m, provided by the GSI (2001), to a
radius of 80 km according to the method by Yamamoto (2002). In all reduc
tions the earth's sphericity is taken into consideration.
Quantitative analysis of gravity anomalies relies heavily on the accuracy of
Bouguer reduction density about which we need a priori information. An error
of 0.1 g / cm 3 in the reduction density, corresponding to an error of nearly 0.42
mgal in Bouguer anomaly for every 100 m, is not a very large error in itself but
an error of 0.1 g / cm3 in the density may have a large effect on the interpreta
tions of Bouguer anomalies (Yamamoto, 1999). However, selection of an
optimum reduction density for the Bouguer and terrain corrections is problem
atic because of the variability in density of the lithologies present. Also we
note that there is not a unique solution for the problem. First, we applied the
82 A. Yamamoto
simple Nettleton's method (Nettleton, 1939) to the central Hokkaido to estimate
an optimum reduction density. The obtained result is 2.68 g / em 3 which is
somewhat a larger estimate. Secondly, we applied the F-H method (Fukao et
al., 1981; Yamamoto et al., 1982) to the same study area to estimate an
appropriate reduction density, and obtained an approximately constant value of
2.61 g/ em3 for mesh size 4' (.-....-6.4 km) to 10' (.-....-16 km). Then, ABlC method
(Murata, 1993) was applied to the same study area and an optimum reduction
density was estimated to be 2.4912 g / em 3. Bouguer anomaly maps produced by
using these optimum densities do not cause considerable changes of amplitudes
or patterns of Bouguer gravity (not shown here). Consequently, we use a
conventional value (2.67 g / em 3) for a gravity reduction density.
Thus Bouguer anomaly values were determined to produce a new gravity
map of the central Hokkaido and adjoining areas. After an overall revision
and reduction for these gravity data, a Bouguer anomaly map, a high-pass
filtered anomaly map, and a first horizontal derivative map of the region were
constructed.
4. Bouguer Anomaly
4.1 Regional-Scale Gravity Anomaly Features
According to the procedures described in the previous sections, we produced
the color-coded new gravity anomaly map in the central Hokkaido and vicinity
(Fig. 3a), where Bouguer gravity is contoured at 2 mgal interval and grid size for
automatic machine-contouring is 0.025 arc-minutes (~40 m). In addition, in
order to emphasize (and isolate) short-wavelength gravity depression and bulge
more fully, Bouguer anomaly fields are high-pass filtered with a cut-off
wavelength of 20 km (Fig. 3b). Thick and red-dotted polygons in Fig. 3 indicate
the rims of the Cenozoic basins defined by Oka (1986). See Fig. 2 for their
sequential numbers (#1.-....-#16). Note that shaded relief of Bouguer gravity in
Fig.3 is generated by illuminating the artificial light from the north-west
direction in order to highlight gravity structures in much more detail.
As shown in Fig. 3a, the Bouguer anomaly field is characterized by two
markedly difference textures divided by a boundary which falls approximately
along the Tokoro Tectonic Line (TTL) of the map area. To the east of TTL,
the Bouguer anomaly field is characterized by the high-amplitude gravity ridge
which is attributable to Mesozoic sequences, whereas to the west, anomaly relief
is much lower with several closed depressions relative to the region to the west
of TTL. Note that these high and low anomaly zones are sharply separated by
44' OO'N
43' 30'N Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 83
km
o 10 20 30
143' OO'E 143' 30'E 144' OO'E
-50 o 50 100
Fig.3a. Bouguer anomaly map of the central Hokkaido and vicinity with a contour
interval of 2 mgal. Assumed density is 2.67 g/cm'- Shaded relief of Bouguer
anomaly is superimposed by illuminatin g the light from the NW direction.
Thick and red-dotted polygons indicate the rims of the Cenozoic basins by Oka
(1986). See Fig. 2 for their sequential numbers (#1 ~#16). See also the caption of
Fig. 2 for the abbreviations and colored faults.
84
44' OON
43' 30N A. Yamamoto
km
o 10 20
-20 -10 o 10 20
Fig. 3b. Same as Fig. 3a, but Bouguer anomalies, with a contour interval of 5 mgal,
are high-pass filtered with a cut-off wavelength of 20 km, where short-wave
length gravity depression and bulge are isolated more fully. 30
IO'N Dense clustering of latest Cenozoic caldera-like basins of central I-Iokkaido 85
r:I!!II!I---"C::;::;:::==~"--"'''C;::==:;::;:;:::;JI.~.-IIIIJ!~;;::;::;:=;;::~~IIJI!.-_C:=~==::JI_-~ 0 H(Sedimcn l3ry rocks)
143· OO'E 143· 30'E 144· OO'E
Fig. 4. Simplifiecl geological map of the central Hokkaido and VICll1!ty Bouguer
anomaly isolines are superimposed for reference with a contour interval of 5
mgal. Geology information is based on Geological Survey of Japan (1995).
Thick and red-dotted polygons indicate the rims of the Cenozoic basins by Oka
(1986). See Fig. 2 for their sequential numbers (#1~#16) . See also the caption of
Fig. 2 for the abbreviations ancl colored faults. o Q3(Scdimcntary rocks) o Q2(Scdimcntary rocks) o Ql(Scdimcntary rocks) o N3(Scdimcntary rocks) o N2(Scdimentary rocks) o Nl(Sedimentary rocks) o PG4(Scdimcntary rocks) o PG3(Scdimentary rocks) o PG2(Sedimentary rocks) o PGl(Scdimcnt.ry rocks) o K2(Marine sedimentary rocks) o Kl·PG2(Scd iment.,)' rocks)
• K2(Mafic volcanic rocks)
• J3-K I (Mafic volctlllic rocks)
• K2(Mafic plutonic rocks) o Ql·H(Volcanicdebris) o N3(Felsic volcanic rocks) o N2(Fclsic volcanic rocks) o Nl(Pelsic volcanic rocks) o Q3-H(Felsic and mafic deposits) o Q2(Felsic and mafic deposits) o Ql(Felsic and mafic deposits) o N3(Felsic and mafic deposits)
• Q3-H(Mafic volcanic rocks) o Q2(Mafic volcanic rocks) o Ql(Mafic volcanic rocks) o N3(Mllfic volcanic rocks) o N2(Mafic volcanic rocks) o Nl(Mafic volcanic rocks) o PG4(Mafic volcanic rocks)
• Nl(Felsic plutonic rocks)
• PG3(Felsic plutonic rocks) o N2(Mafic plutonic rocks)
• Nl(Mafic plutonic rocks) II PG3(Mafic plutonlc rocks)
KI-PGI(Mclamorphic rocks)
CENOZOtC
H: Holocene
Q:QUtllcnury
N: Neogene
PO: Pnleosene
MESOZOIC
K:Crel:lCeous
J: Jurassic
TR:Triassic
(I:Eariy. 2:Middle. 3:Late)
86 A. Yamamoto
TTL running in NWN -SES direction. The Bouguer anomaly ridge to the east
of TTL dominates extensively in the Mesozoic sequences around the eastern
part of the map area whose easternmost and westernmost boundaries are almost
coincident with the Abashiri Tectonic Line (A TL) and TTL, respectively.
Apparent gravity anomaly depressions in ellipsoidal shape with almost NE-SW
major axis of 15~20 km sharply penetrate this gravity ridge to the eastern edge
of the map area (Fig. 3a).
In contrast, the gravity low field to the west of TTL is divided into two
parts: one occurs around the Tokachi Tectonic Basin (TTB) which is mor
phologically bordered around the portion passing through Rikubetsu (RBT), Mt.
Kitoushi (KIT), and Mt. Higashi-nupukaushinupuri (HNP), the other consists
of the vast gravity depression area situated north of TTB. These gravity lows,
extending northward of TTB, take their minimum (about -46 mgal) near Mt.
Asahidake around the Daisetsu and Tokachi volcanic areas (Fig.3a). Un
equivocally, these remarkable lows form a cluster of closed depression in a
linear dimension of about 10 km in diameter (Fig.3b). These salient features
are strengthened by a relief-shaded Bouguer image in Figs.3a and 3b. It is
quite remarkable that Bouguer anomaly contours of these marked lows tend to
become dense at a rim of each closed depression, which is discussed later.
4.2 Gravity Anomaly around the Basin Area
As shown in Figs. 1, 2 and 3, most of these Bouguer lows also coincide with
the distributions of late Cenozoic basins. This is the most surprising feature
which yields constraints on the geometry of underground structures of the
basins. Although the basins #1 and #5 (Figs. 2 and 3) do not show a closed
Bouguer depression where distinct basin topography does not develop (Fig. 5),
trend-reduced Bouguer map shows a low anomaly whose anomaly relief is more
gentle. As shown in Fig. 3b, several gravity high-low transitions can be obser
ved inside the basin #1, while around the basin #5 occurs a gentle gravity low.
We found a good correlation between closed low anomalies and the location of
the Cenozoic basins #6~#8. Particularly, the basin #6, which is known as "the
Shirataki Basin", displays a surprising coincidence between its rim and the
location of the steep gravity change (maximum in horizontal gravity gradient,
see Fig. 6). Similar coincidence appears in the basin #7 while Bouguer anomaly
changes around the basin rim are not so steep and Bouguer gravity decreases
gently, not abruptly, toward the basin center, reaching a minimum near the
center. Also the basin #8 appears to have a closed depression of gravity field.
44· OO'N
43· 30'N Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 87
km
o 10 20 30
144· OO'E
"IIIIlI ... ".... ..... ~~ ...,.......,. ..... ....iii-II-............... .... _-V Topography(m) 143· OO'E 143· 30'E
o 500 1000 1500
Fig. 5. Digital topography and known faults in the central part of I-Iokkaido, il·
luminated by the light from the NW direction. Bouguer anomalies are contoured
at 5 mgal interval for reference. See the caption of Fig. 2 for the abbreviations
and colored faults.
88 A. Yamamoto
Gravity field of the Cenozoic basins #2~#4 is also characterized by a sharp and
closed dent whose relative amplitude amounts to be 10~20 mgal (Figs. 2 and 3).
In particular, the Cenozoic basins #4 forms an almost ring-shaped Bouguer low
with characteristically steep rim which indicates nearly the same location of the
basin perimeter. Similarly, the basin #3, known as the Tokachi-Mitsumata
Basin, is clearly outlined by the gravity depression associated with a steep
gradient rim (see Fig. 6), where generalized topography around the basin also
shows that the edifice is almost ring-shaped (Fig. 5). Interestingly, low Bouguer
anomaly of the basin #2 seems to be stressed in trend-reduced Bouguer gravity
as shown in Fig. 3b.
Ishii (2001) and Ishii et al. (2002) argued about the characteristic features
of the Tokachi-Mitsumata Basin and pointed out that the basin shows distinct
low gravity anomalies using the same gravity dataset in this study. Finally,
they concluded that the Tokachi-Mitsumata Basin, formed in late Cenozoic age,
is a circular-shaped caldera from geological evidences. As shown in Figs. 2, 3
and 4, the "Tokachi-Mitsumata Caldera" morphologically has an obvious ring
shaped rim which seems to control the Bouguer contours particularly in the
northern side of the caldera. The southward extension of this low anomaly
continues 5 or 6 km south of this caldera and seems to adjoin the closed Bouguer
low of the basin #2 whose gravity field has rather moderate gradient to the
center. Gravity anomaly features around the Cenozoic basins #9~#13 are
similarly expressed by an oval-shaped depression with a relative amplitude of
more than 10 mgal. Particularly, in the basins #10 and #13, a good coincidence
occurs between their rims and the location of a steep gravity change. On the
contrary, the Cenozoic basin #9 largely indicates gravity swelling which is
inversely correlated with the topographic depression associated with the basin,
although the anticipated gravity feature is a Bouguer low. While the Cenozoic
basins #11 and #12 does not correlate well with low gravity fields. It should be
noted that immediately about 5 km west of the basin #12 occurs another
remarkable gravity dent which is not correlated with any known basins. This
N -S elongated gravity low, whose center is situated about 12 km ESE of
Maruseppu (MSP) between the calderas #11 and #13, has almost the same
diameter as the basin #13 with a relative amplitude of about 15 mgal as shown
in Fig. 3b. We find the conspicuous gravity lows in the Cenozoic basins that are
well-developed over the Mesozoic regions around the eastern part of the map
area. These low anomalies well correlate with the distributions of the Cen
ozoic basins #14~#16, each of which forms an oval-shaped topographic depres
sion controlling Bouguer contours and runs almost in NE-SW direction. It is
Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 89
worth noting that gravity gradients outside each basin but sloping toward it are
clearly found to be rather gentle than those inside, which is interpreted as
evidence of a low-density mass located below the basin. As described above,
most of the Cenozoic basins indicate a good correlation with the closed depres·
sions of gravity field, which strongly suggests that these basins possess caldera
like morphology with a piston-shaped form. This feature will be evidenced
more clearly by the steep gravity gradients over nearly ring-shaped perimeter
of the basins.
In addition, we point out several more gravity depressions which would not
be correlated with any of known Cenozoic basins. The existence of these
gravity lows suggests that there would be a hidden basin or caldera below the
gravity depression areas. One of the most distinguishable hidden structures is
located at the central part of the map area, entirely surrounded on three sides
by the Cenozoic basins #4 and #9 (on the south side), #6~#8 (on the west side),
and #10~#13 (on the east side). Maruseppu (MSP) is located near the northern
most part of this notable depression. This area shows a large-scale gravity
depression about 10 km (EW) by 25 km (NS) and consists of several closed
minima, each of which indicates a sharp gravity change around its rim.
Another hidden structure can be found 5 km south-west of Rubeshibe (RSB)
which is characterized by a circular-shaped gravity low with a relative ampli
tude of more than 10 mgal. The Bouguer anomaly inside or at the rim of this
low, however, changes rather moderately and gradually decreases toward its
center.
4.3 Horizontal Gradient
As became clear in the previous discussion, it is notable that most of the
gravity depressions corresponding to the Cenozoic basins indicate a steep
gravity gradient (horizontal derivative) along the basin rims. As might be
expected easily, large gradient values along a basin rim infer that the walls of
the basin dip with high angle (vertically of inward). The most straightforward
interpretation for this feature is that the basins may have piston-shaped, not
funnel-shaped, subsurface structure located beneath the basin, where the wall
dips with high-angle.
Recent gravimetric analyses have shown a useful approach using horizontal
derivative (horizontal component of gravity gradient) in order to help to (1)
delineate remarkable zones which indicate abrupt gravity changes (steep gravity
gradient) (e.g., Yamamoto et at., 1986), and (2) highlight gravity structures
44· OO'N
43· 30'N (a) km r=- r=- r=- I o 10 20 30
, )
.>--~.
/"\-\
~
'-..
'I, I,. .1S0
'~ ,r.f<., ("" ,-/ '...../ 44· OO'N
/
,.) j(MW
\
\
43· 30'N
.. "'"
143·00'E 143·30'E 144· OO'E
~ .. Ii ~ mgallkm
o 2 3 4 6 8 9 10 (b) km
o 10 20 30
vf'l\
~ , )
?---'\~
.;;-
~ , \ .TSO f',,>J:'
<.
(
V
\ \
<'/
.. j(MW
143·00'E 143· 30'E
~C~~~~~~==~ ~ ~ mgal/km
o 6 7 8 10 2 3 4 144· OO'E
Fig. 6. (a) Bouguer anomaly gradient map around the central Hokkaido for gradient values ranging from 0 to 10 mgal/km. Yellow-coded ridge portions
delineate locus of large inflection points. (b) Same as (a), but the perimeter s of the Cenozoic basins by Oka (1986) are superimposed by thick and red
dotted polygons for reference . Their sequential numbers (#1~#16) are shown in Fig. 2. Note a remarkable coincidence between the perimeters of several
basins and maximum inflection of gradient fields, See the caption of Fig. 2 for the abbreviations and colored faults,
Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 91
laterally varying along the specific direction (e.g., Yamamoto, 2003, 2004).
Accordingly, in order to examine the quantitative aspects of the gradient fields
more closely, we take horizontal derivative of the 2-D Bouguer anomaly,
particularly to investigate the variations of subsurface structure with special
reference to the Cenozoic basins or caldera of the central Hokkaido. In this
study, according to the past gravity works (e.g., Cordell, 1979; Cordell et at.,
1985; Yamamoto et at., 1986), horizontal derivative of 2-D Bouguer gravity
field is defined by
[( og/ox)2 + (og/oy )2]1/2,
where g(x, y) is Bouguer gravity field, x and y indicate longitudinal and
latitudinal directions, respectively. Units are milligals per kilometer. Note
that the directional properties are completely lost in the above formula. It
should be also noted that ridges in the amplitude of horizontal gradient field
generally delineate the traces of gravity inflection points, by inference, the
traces of basin-border or caldera-bounding faults. Using the above formula,
Bouguer anomaly gradient map (Fig. 6) is produced around the central
Hokkaido for gradient values ranging from 0 to 10 mgal/km. The location of
the Cenozoic basins is superimposed in Fig. 6b for reference. Notice that high
gradient zones are expressed distinctly by yellow colors in Fig. 6. Several
major high-gradient zones can be found to be nearly closed in a ring-shaped or
oval-shaped form in the central part of the map area. As shown in Fig.6b,
some of these closed extremes (maximum trace in horizontal gravity gradient)
notably accord with the rim of the Cenozoic basins. Particularly in the basins
#2~#4, #6, #10 and #16 a remarkable coincidence between their rims and the
location of a steep gravity change can be found, whose maximum value reaches
more than 13 mgal/km around the basins #4 and #10. The Bouguer contours of
these basins are most likely to be those of the thickness contours of the basin fill,
which places good constraints on the subsurface structure below the basin.
Thus, it is quite significant that the boundary of the Cenozoic basins (calderas)
seems to be defined as coinciding with the maximum trace in horizontal gradient
of Bouguer anomaly gravity field (or vice versa), which may be interpreted as
the result of several collapses associated with volcanic activities. These facts
also imply that caldera-like collapse structure filled with a low-density material
develops well beneath the Cenozoic basins of the central Hokkaido.
92 A. Yamamoto
5. Discussion
The average density of the basin fill depends on many poorly-known factors
and can only be estimated with considerable uncertainty. Much of the porous
sediments probably have a density value of about 2.2 g / em 3 or smaller. Some
of the volcanic units, however, have physical densities which equal or exceed the
density of the enclosing crystalline textures. Pumious or scoriacious mate
rials, in particular, may have extreme anomalous density value considerably less
than 2.0 g / em 3. So if we assume the average density of the basin fill of the
"Tokachi-Mitsumata Caldera" (the caldera #3, see Fig. 2) to be 2.2g/em3,
yielding a density contrast of 0.47 g/em3 in this study, simple calculation shows
that an apparent maximum thickness (depth) of 750 m~1,000 m, which corre
sponds to a relative amplitude of 15 and 20 mgal. In this case several isolated
closure of Bouguer contours are found within the caldera with a relative
amplitude of about 2 or 3 mgal, indicating that the "Tokachi-Mitsumata
Caldera" has a fiat-bottom structure bounded by a steep (or high-angle) wall
corresponding to the caldera-bounding rim. This feature suggests that the
caldera may have originated in volcanic depression which subsided along high
angle normal fault, although generalized topographic features around this area
are poorly resolved due to collapse of the caldera walls.
Similarly in the "Tokachi-Mitsumata Caldera", Bouguer anomalies of
other basins, which display a surprising coincidence between their rims and the
location of a steep gravity change, have a relative amplitude of about 10~20
mgal. Accordingly, this feature suggests that most of the Cenozoic basin of the
central Hokkaido may possess a low-density basin fill with an apparent thick
ness (depth) of 500 m~1,000 m, assuming a density contrast of 0.4~0.5 g/ em3•
Although perturbations of the density contrasts and geometric configurations
would improve the model fit, little new information would result without first
obtaining a gravity dataset with significantly smaller station spacing to accu
rately define gravity anomaly. It is worth noting that the subsurface structure
below these basins can be successfully constrained by Bouguer gravity anomaly,
though volcanic collapse and infilling history of the Cenozoic basins in the
central Hokkaido may be complicated and it is beyond the scope of this paper
to deal with that matter.
6. Summary
We produced a new gravity anomaly map, as well as a Bouguer gradient
Dense clustering of latest Cenozoic caldera-like basins of central Hokkaido 93
map, of the central Hokkaido which enabled delineation of the main features as
follows : (1) To the east of the Tokoro Tectonic Line (TTL), the gravity field
is characterized by high-amplitude anomalies which outline the predominant
Mesozoic sequences. These gravity ridges are bounded by steep gradients in
the westernmost part. Whereas to the west of the TTL, anomaly relief is much
lower, associated with several closed dents which correlate well with the
distributions of the Cenozoic basins, (2) Several remarkable gravity lows form
a cluster of closed depression with a diameter of about 10 km, and Bouguer
anomaly contours of these marked lows tend to become dense at a rim of each
closed depression, which suggests that the walls of the basin dip with high angle
(vertically of inward). Thus, a gratifying aspect of this study has been emphas·
ized by the remarkable facts that the boundary of the Cenozoic basins (calderas)
seems to be well defined as coinciding with the maximum inflection in horizontal
gradient of Bouguer anomaly gravity field, whereas Bouguer gradient fields
inside most of the basins show considerable smaller values which implies a flat
bottom. These facts corroborates that the basins possess a piston-shaped, not
funnel-shaped, underground structure which may be the result of several distinct
collapses, as long as the gradient field shows maximum inflection along their
rims.
In addition, several more gravity depressions, none of which would be
obviously correlated with known Cenozoic basins nor morphological features,
are clearly observed. Most of these distinguishable hidden structures gather at
the central part of the map area, entirely surrounded on three sides (east, west
and south) by the Cenozoic basins. The existence of these gravity lows strongly
suggests that there would be an invisible basin or caldera below the gravity
depression areas.
Assuming the average density of the caldera fill to be 2.2 g / cm3 in one of the
typical calderas (the "Tokachi-Mitsumata Caldera") in the central Hokkaido,
which yields a density contrast of 0.47 g/cm3, an apparent maximum thickness
of the caldera fill amounts to be approximately 750 m~ 1,000 m, which corre·
sponds to a relative amplitude of 15 and 20 mgal, respectively.
We conclude that digital processing (horizontal derivatives, high-pass
filtering and relief-shaded imaging) proved useful in establishing the conceptual
underground model of the basins, which would be supported by 2-D or 3-D
numerical modeling.
94 A. Yamamoto
Acknowledgements
I am grateful to Haruyoshi Ishikawa of the Institute of Seismology and
Volcanology, Hokkaido University, for his help. Discussions with Mitsuhiro
Nakagawa and Eiichi Ishii of Hokkaido University were helpful. The produc
tion of the figures was facilitated by the freely available GMT software (Wessel
and Smith, 1995; Available as http://www.soest.hawaii.edu/gmt/ ).
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Journal of Mineralogical and Petrological Sciences,Volume 106, page 255-260, 2011 LETTER Petrology and geochemistry of chromian spinel-bearing serpentinites in the Hida Marginal Belt (Ise area, Japan): characteristics of their protoliths Mohamed ZAKI KHEDR*** and Shoji ARAI *Department of Earth Sciences, Graduate School of Natural Science and Technology, Kanazawa University, Kakuma, Kanazawa 920-1192, Japan *Department of Geology, Faculty of Science, Kafrelsheikh University, Egypt The Ise area is located in the western part of the Hida Marginal Belt (central Japan), which includes several spo- radic exposures of ultramafic rocks, sometimes forming a serpentinite mélange of the Paleozoic age. Ultramafic rocks in the Ise area, enveloped by Paleozoic-Mesozoic sedimentary rocks, are completely serpentinized; how- ever, the abundance of bastite after orthopyroxene suggests harzburgite protoliths. The bastite- and mesh-tex- tured serpentines are distinguished from each other in AlzO3, CrzOs and NiO contents. The bastite-textured ser- pentine is high in Al2O (up to 4.0 wt%) and Cr2O (up to 1.2 wt%), but low in NiO (<0.3 wt%) relative to the mesh-textured one. Relic chromian spinel, vermicular in shape, shows an inter-grain chemical homogeneity tio] from 0.38 to 0.51 and low YFe {=[Fe?+/(Cr + Al + Fe?+) atomic ratio, <0.03]}, similar to chromian spinel in Kotaki and Oeyama ultramafic masses. It is also similar in chemistry to spinels in forearc and abyssal perido- tites, suggesting two possibilities for the derivation of the Ise serpentinite's protoliths. The degree of melting us- ing Cr# versus TiO2 of chromian spinel is ~ 20-25%, which is in accordance with the harzburgite protoliths ob- tained by whole-rock chemistry models. We found that Cr, Al, Ni and Ca, preserved in bastite and mesh- textured serpentine, are conservative during serpentinization, confirmed from the similarity in the whole-rock A1 and Ca of the Ise serpentinites to Horoman harzburgites. The occurrence of dolomite can stabilize the Ca and limited its mobility to escape outside the serpentinites. The harzburgite protoliths were possibly serpen- tinized by slab-derived fluids at a shallow depth relative to the Happo-O'ne serpentinites during the exhumation process. Keywords: Chromian spinel, Bulk chemistry, Serpentinites, Ise, Hida Marginal Belt, Japan INTRODUCTION serpentinization can obliterate the textures and magmatic signatures of most peridotites as in the Happo-O'ne peri- The alpine-type ultramafic complexes have been em- dotites (e.g., Nozaka, 2005; Khedr and Arai, 2010). We placed into the Paleozoic sediments and surrounded by found that some Ise serpentinites have preserved pseudo- crystalline schists in western Japan (Arai, 1980). Ultra- morphic textures (e.g., bastite and mesh) after primary mafic rocks of the Ise area, central Japan, are completely minerals. These textures and intact primary chromian spi- serpentinized, but their chromian spinels are partly intact. nel can decipher the features of the Ise serpentinite's pro- The Cr# and AlzOs values of chromian spinel are used to toliths. Information from the ultramafic rocks in the Ise constrain the nature of the mantle peridotites and the de- area supplementary to that from the Happo-O'ne ultra- gree of partial melting as treated in the literature (Dick mafic complex in the Hida Marginal Belt (HMB) help us and Bullen, 1984; Bonatti and Michael, 1989; Arai, 1994). to know the tectonic evolution of HMB. We try to provide Recently, Niu (2004) stated that serpentinization has not constraints on element behaviors during serpentinization obliterated the magmatic signatures in the whole-rock processes. compositions for abyssal peridotites. It is well known that doi:10.2465/jmps.110606 M. Z. Khedr, khedrzm @ yahoo.com Corresponding author 256 M. Zaki Khedr and S. Arai GEOLOGIC SETTING ay 6 The Hida Marginal Belt in central Japan includes several masses with sporadic exposures of ultramafic rocks, e.g.. the Ise area, Fukui Prefecture. It includes mafic and ultra- mafic rocks associated with high P/T metamorphic rocks of the Renge metamorphic belt, which represents a sub- duction zone complex of Paleozoic age in the southwest- ern Japan (e.g., Nakamizu et al., 1989). The large masses of ultramafic rocks (1.5 Km long-1.0 km across) are em- bedded in a fine-grained matrix (Paleozoic-Mesozoic sediments), forming a serpentinite-matrix melange in the Ise area (Fig. 1: Figure 1 is available online http://joi.jlc. jst.go.jp/JST.JSTAGE/jmps/110606.). This mélange con- sists of high-P/T Renge metamorphic rocks (283-338 Ma; e) T Kunugiza et al., 2004), serpentinites, gabbros, amphibo- lites, greenschists and blueschists that are in direct contact with carbonate and clastic sedimentary rocks (Miyakawa, 1982; Yokoyama, 1985; Kurihara, 2003; Tsujimori et al., 2006). The Ise area is bounded by a Jurassic accretionary 0.5mm complex (Mino-Tanba Belt) on the south side, and Paleo- zoic sediments and metamorphic rocks on the north and Figure 2. Photomicrographs of the serpentinites from the Ise area. (a) The pseudomorphic textures, such as mesh serpentine re- west sides (Fig. 1). Ultramafic masses in the Ise area are placements of olivine and bastite serpentine replacements of or- dissected by faults and bounded by metamorphosed Pa- thopyroxene (No. Es-27). Crossed-polarized light. (b) Dolomite leozoic rocks (Ise crystalline schist) on the south side. granules scattered in the center of mesh texture (No. Es-21). They are completely serpentinized and enveloped by Pa- Crossed-polarized light. (c) The blades and flakes of antigorite, leozoic-Mesozoic sedimentary (clastic and carbonates) forming interpenetrating textures (No. Es-7). Crossed-polarized light. (d) The chromian spinel, vermicular in shape, enclosed in rocks (Kurihara, 2003). The serpentinites are mainly mas- the pseudomorphic orthopyroxene (bastite) that surrounded with sive and enriched with carbonates in some parts along the dolomite (No. Es-20). Plane-polarized light. (e) Deep red col- Ise River. ored chromian spinel preserved in serpentinites (No. Es-20). Plane-polarized light. (f) Back scattered electron image of chro- PETROGRAPHY mian spinel in e panel. Abbreviations; Dol, dolomite: Spl, prima- ry chromian spinel: Serp, serpentines. The ultramafic rocks in the Ise area are completely ser- pentinized. They were divided into two serpentinite gro- acteristics reflect equigranular to porphyroclastic textures ups: one around Nigure River (Es-1 to Es-12) and another of the protoliths, which are possible harzburgites. Antig- one around the Ise River (Es-13 to Es-33). The serpenti- orite occurs mainly as blade or flame-like crystals, form- nites around the Ise River include subordinate amount of ing interpenetrating textures (Fig. 2c). Carbonates occur carbonate (<5%), chromian spinel (<2%), magnetite mainly as granules (0.2-0.3 mm) in serpentine mesh cen- (<5.5%), chlorite (<1%), and sulphides (<1%). The bas- ters and along relic cleavages in bastites after orthopyrox- tite- (15-23%) and mesh-textured serpentines after or- ene (Figs. 2b and 2d). Chlorite occurs as fakes around thopyroxene and olivine, respectively, are dominant in chromian spinels. Primary chromian spinel, deep red in most samples. Whereas serpentinites around Nigure River thin section (Figs. 2e and 2f), is vermicular in shape (up show interpenetrating texture (antigorite) with a minor to 2.0 mm across), and sometimes enclosed by orthopy- trace of bastite texture, and are free of chromian spinel roxene pseudomorphs (Fig. 2d). Some chromian spinel and chlorite. Serpentine minerals, when specified, were grains are altered to magnetite and/or feritchromite along identified by a laser Raman spectrometer. Bastite after or- their cracks. Magnetite occurs mainly as disseminated thopyroxene, 1.0-6.5 mm in length, occurs as coarse sub- grains. It is sometimes associated with sulfides that are hedral prismatic crystals and consists of chrysolit/lizardite mainly millerite, siegenite, and violarite with subordinate (Figs. 2a a nd 2d). The mesh-textured part comprises Co-gersdorffite (NiAsS). equigranular (1.2-1.1 mm) aggregates of serpentine (anti- gorite or lizardite) (Figs. 2a and 2b). These textural char- Petrology and geochemistry of chromian spinel-bearing serpentinites 257 CHEMISTRY 0.9 wt%). The bastite-textured serpentine is high in AlO3 OIN u! moI 1nq “(%m 7'I-7'0) O1 pue (%m +-7'1) The Ise massive serpentinites with relic porphyroclastic to (<0.3 wt%) relative to the mesh-textured one (<1.0 wt% equigranular textures were selected for the whole-rock Al2O3, <0.2 wt% Cr2O3, and up to 0.9 wt% NiO) (Fig. 4: analysis. Their powders were heated up to 1000 °C before Figure 4 is available online http:/joi.jlc.jst.go.jp/JST.JSTAGE/ preparing fused disks for major-element analysis that was jmps/110606.). Both types of serpentine have low CaO carried out by XRF (System-3270, Rigaku) at 50 kV ac- (<0.2 wt%), TiO2 (<0.06 wt%) and Na2O (<0.03 wt%) celerating voltage and 20 mA beam current at Kanazawa (Table 2). This relationship between the texture and the University (Table 1: Table 1 is available online http://joi. composition of serpentine minerals shows some similarity jlc.jst.go.jp/JST.JSTAGE/jmps/110606). The JB-2 (basalt) to the Guatemala forearc serpentinites (Kodolanyi and reference sample obtained from the Geological Survey of Pettke, 2011). The examined bastite shares in the same Al Japan (GSJ) was used for calibration. The accuracy of and Cr values with primary orthopyroxene in Oeyama measurements based on JB-2 reference sample (Imai et harzburgites (Arai, 1975, 1980; Kurokawa, 1985) (Fig. 4). al., 1995) is acceptable and agrees with previous values in Chlorite is mainly penninite in composition, and shows a the literature (Table 1) at the 92-100% confidence level. narrow range of Mg# (0.94-0.95) (Table 2). Dolomite The major-element contents (Tables 2 and 3: Tables 2 and (Mg#, 0.95) and magnesite (Mg#, 0.88) the main carbon- 3 are available online http://joi.jlc.jst.go.jp/JST.JSTAGE/ ate phases are high in Mg# (Table 2). jmps/110606.) of silicates and spinels in the Ise serpenti- The chromian spinel cores show an inter-grain che- nites were analyzed by an electron microprobe (JXA- mical homogeneity, and has an average composition of 8800, JEOL) at Kanazawa University. Accelerating volt- 31.5 wt% Al2O3, 14.0 wt% MgO, 36.5 wt% Cr2O3, 17.5 age, beam current, and beam diameter for the analyses wt% FeO*, and 0.03 wt% TiO2 (Table 3). It has a narrow were 20 kV, 20 nA, and 3 μm, respectively. Ycr, Yai and range of Cr# [= Cr/(Cr + Al) atomic ratio] from 0.38 to Yre are the atomic ratios of Cr, Al and Fe?+ in spinel, re- 0.51 (0.44 on average) and low Yre {=[Fe+/(Cr + Al + spectively, to trivalent cations, (Cr + Al + Fe3+). Fe?+) atomic ratio, <0.03]}. The chromian spinel, which is sometimes altered to ferritchromite (Cr#, up to 0.98 ） at WHOLE-ROCK CHEMISTRY the margin, shows chemical zoning in which AlO3, MgO, and CrzO3 decrease severely toward the rim (enriched in The Ise serpentinites share the same composition of AlO3 FeO). The Ise harzburgite lies in the overlap zone between (0.33-1.13 wt%) and CaO (0.23-0.81wt%) with the Horo- forearc (Ishii et al., 1992) and abyssal peridotites (Dick man harzburgites (Takazawa et al., 2000), the Izu-Bonin- and Bullen, 1984) in spinel chemistry (Fig. 5); it is similar Mariana (IBM) forearc peridotites (Parkinson and Pearce. in spinel chemistry to mantle harzburgites in the Oeyama 1998) and the Happo-O'ne metaperidotites (Khedr and ophiolite (Arai, 1980; Kurokawa, 1985) (Fig. 5). The ex- Arai, 2009, 2010; Khedr et al., 2010) (Figs. 3a and 3b: amined chromian spinel seems to have intermediate spinel Figure 3 is available online http:/joi.jlc.jst.go.jp/JST.JSTAGE/ composition between Oeyama group (Cr#= ~ 35) and jmps/110606.). They are similar in major-element compo- Tari-Ashidachi group (Cr#= ~ 50) of lherzolites-harzbur- sitions to Happo-O'ne metamorphosed peridotites, except gites from the Kotaki area (Machi and Ishiwatari, 2010) their high in SiO2. Total Fe as Fe2O3 (7.9-10.8 wt%) is (Table 3; Fig. 5b). higher than that of the primitive mantle (PM; Niu, 1997) as in the Happo-O'ne metaperidotites. Their whole-rock DISCUSION Mg# [=Mg/(Mg + Fe2+) atomic ratios)] ranges from 0.89 to 0.92, like that of the Happo-O'ne metaperidotites Element behavior during serpentinization of Ise peri- (Khedr et al., 2010), reflecting the depleted nature of their dotites protoliths. Calculation based on the bulk-composition calculation (Niu, 1997) yields low clinopyroxene modal The abundance of bastite-textured serpentines (15-23 amounts% (1.2-3.8%), supporting the harzburgite proto- - ds sisns xno re ( ro lith judged from the petrography (Table 1). burgites are the protoliths (Figs. 2a and 2d). So, we com- pared the Ise serpentinites in whole-rock chemistry with MINERAL CHEMISTRY the known Horoman harzburgites (Takazawa et al., 2000). The low Al2O3 (<1.2 wt%) and CaO (<0.82 wt%) contents Serpentines, lizardite/chrysotile and antigorite, show wide are compared to those of Horoman harzburgites (Figs. 3a range of Mg# (0.89-0.97), Al2Os (0.0-3 wt%), Cr2Os (0.0- and 3b), being equivalent to low modal clinopyroxene 1.8 wt%), FeO as total Fe (0.9-7.0 wt%) and NiO (0.01- (<4%; Table 1). This similarity in Al and Ca contents with 258 M. Zaki Khedr and S. Arai (a) Cr the dolomite (Figs. 2b and 2d) can be a reservoir of Ca re- Ise Serpentinites leased from peridotitic minerals and limited the mobility O Cr-spinel Fe-ch to Mt of Ca. This is consistent with the limited mobility of Ca, Happo-O'neDunites mobile only in a small scale in serpentinite samples, dur- Cr-spinel and its alteration ing serpentinization of abyssal peridotites (Niu, 2004). Also, the mobility of Ca (dissolved from diopside) in flu- Omi serpe ids during serpentinization processes is controlled by sili- ntinites ca activities and fluid conditions (pH, oxygen fugacity) (Frost and Beard, 2007). We suggest that the whole-rock Lower Al and Ca contents have been preserved in the Ise serpen- amphibolite facies tinites. The whole-rock Fe and Si contents are higher than those of PM (Niu, 1997), possibly reflecting Si- and Fe- Greenschist enrichment during subduction-related metasomatism prior facies to or during serpentinization processes. There is possibly AI Fe3 cryptic metasomatism for Si. 1.0 (q) Omi serpe The bastite-textured serpentine is high in AlO3 and ntinites Cr2O3, but low in CaO, TiO2 and NazO (Fig. 4 and Table Alpine-type! 0.8 peridotites? 2), suggesting basically being inherited from the precur- sory orthopyroxene composition. Its similarity in Al and Forearc peridotites Cr values to primary orthopyroxene in Oeyama harzbur- 0.6 gites (Arai, 1975, 1980; Kurokawa, 1985) is in agreement # with inheriting of these elements from the precursor Opx. Harz(Oeya 0.4 Ophiolite) The low Ca content of partly bastite is due to the low Ca Lhrz/Harz character of the precursor orthopyroxene but partly due to (Kotaki area) the consumption of Ca to form dolomite within bastite 0.2 (Figs. 2b and 2d). We suggest that Al, Cr and Ca might Abyssal peridotites have been immobile during serpentinization of the Ise- harzburgite protoliths. 1.0 0.8 0.6 0.4 Mg# Tectonic implications for the protoliths of Ise serpenti- Figure 5. Chemical characteristics of chromian spinel in the Ise serpentinites. (a) Al-Cr-Fe3+ trivalent cation diagrams. (b) Rela- nites tionship between Cr# and Mg# of chromian spinel. The exam- ined spinel lies in the overlap zone between forearc and abyssal The composition of chromian spinel is used to constrain peridotites. Chromian spinel and its alterations in the Happo- the nature of mantle peridotites and the degree of partial O'ne dunites are obtained from Khedr and Arai (2010). Fields of chromian spinel in forearc peridotites (Ishii et al., 1992), Alpine- melting (Dick and Bullen, 1984; Bonatti and Michael, 1989). The relic chromian spinel (Cr#= 0.38-0.51) in the type as well as abyssal peridotites (Dick and Bullen, 1984), lher- zolites-harzburgites from the Kotaki area (Machi and Ishiwatari, Ise serpentinites is similar to spinels in Oeyama harzbur- 2010), mantle harzburgites in the Oeyama ophiolite (Arai, 1980; gites from the Tari-Misaka complex (spinel Cr#=0.5; Kurokawa, 1985) and Omi serpentinites (Tsujimori, 2004) are Arai, 1980), confirming the harzburgite protoliths as in- shown for comparison. The grey arrowed line indicates a trend of ferred from the whole-rock chemisty and the texture. It spinel composition change during retrograde transition, lower amphibolite- to greenschist-facies (Muntener et al., 2000). Ab- has intermediate composition between spinels (Cr#= breviations; Fe-ch, ferritchromite: Mt, magnetite: Harz, harzbur- 35) in Oeyama group and Tari-Ashidachi group (Cr#= ~ gites: Lherz, lherzolites. 50) from the Kotaki area (Machi and Ishiwatari, 2010) (Fig. 5b), possibly providing a missing link between the two groups, which may have been originally continuous. the Horoman harzburgites possibly indicates that these el- In terms of the spinel chemistry, the Ise harzburgite is ements are conservative during serpentinization. This is similar to some forearc peridotites (Fig. 5). However, it consistent with the immobility of A1 during serpentiniza- has been also derived from low-T abyssal peridotites be- tion (e.g., Bonatti and Michael, 1989; Niu, 2004). It is cause the Mg# of spinel in peridotites is lowered due to well known, however, that Ca is subtracted from perido- cooling along with hydration (e.g., Okamura et al., 2006; tites during serpentinization to change gabbroic rock to Hirauchi et al., 2008). Considering the tectonic setting of rodingites (e.g., Frost and Beard, 2007). We suggest that the HMB peridotites (cf., Khedr and Arai, 2010), the Ise Petrology and geochemistry of chromian spinel-bearing serpentinites 259 serpentinite's protoliths were possibly derived from the have been formed at low temperatures (T < 600 °C). This mantle wedge at a sub-arc setting, like Oeyama perido- result is consistent with the common association of antig- tites (Matsumoto et al., 1995; Arai, 1997). Based on the orite with dolomite and magnesite in the Ise serpentinites, Cr#-TiO2 relation of spinel (Fig. 6b), the Ise harzburgite reflecting temperature from 450 to 580 °C based on a pet- possibly represents residues after ~ 20-25% melting rogenetic grid for carbonate-bearing hydrous ultramafic (Pearce et al., 2000). This is supported by using Al and Ca rocks (Will et al., 1990). The protoliths of Ise serpenti- of the whole-rock chemistry (e.g., Ishiwatari, 1985; Niu, nites may have been hydrated by slab-derived fluids at a 1997) (Fig. 3). The spinel in the Ise harzburgite is en- shallow depth relative to the Happo-O'ne serpentinites riched in Al relative to that in the Happo-O'ne dunites during the exhumation process, forming chromian spinel- (Fig. 5a). It is possibly similar to spinel in the Happo- bearing serpentinites at amphibolite facies. O'ne metaharzburgite, which has been altered to form low-T minerals (chlorite and antigorite) (Khedr and Arai, ACKNOWLEDGMENTS 2010). The compilation of spinel chemistry from Happo- O'ne and Ise ultramfic rocks are used to decipher origins We are grateful to Mr. N. Akizawa, Mr. M. Negishi, Mr. and characheristics of Paleozoic mantle wedge-peridotites M. Miura, Ms. K. Abou-kebir and Dr. B. Payot, for their in HMB. assistance during the field work in the Ise area. We are also indebted to Dr. T. Mizukami and Mr. M. Miura for The antigorite interpenetrating texture (Fig. 2c) is formed at the stability field of antigorite (250-600 °C; Ev- their help during a laser Raman analysis. We are grateful ans, 2004). The Ise serpentinites of mineral assemblage to Drs. T. Tsujimori and Y. Ohara for their beneficial com- chlorite + ferritchromite + antigorite are stable at the am- ments. We thank Prof. Dr. A. Ishiwatari for his support phibolite facies (e.g., Evans and Frost, 1975). The exam- and editorial handling of this paper. ined antigorite serpentinites, free of bastite and mesh tex- DEPOSITORY AND SUPPLEMENTARY tures (samples around Nigure River), have been possibly MATERIALS formed at T (400-600 °C) in the early stage, like the Hap- po-O'ne serpentinites that are composed mainly of antig- Figures 1, 3, 4 and Tables 1, 2, 3 are available online orite (Nozaka, 2005, Khedr and Arai, 2010; Khedr et al., http://joi.jlc.jst.go.jp/JST.JSTAGE/jmps/110606. 2010). The studied mesh and bastite textures, being absent REFERENCES in the Happo-O'ne serpentinites, are composed mainly of chrysotile and lizardite, which were formed at more surfi- Arai, S. (1975) Contact metamorphosed dunite-harzburgite com- cial environments (T < 400 °C) (Evans, 2004). The ser- plex in the Chugoku district, western Japan. Contributions to pentinites around Ise River, composed mainly of lizardite Mineralogy and Petrology, 52, 1-16. + chrysotile ± antigorite, have been formed at low T Arai, S. (1980) Dunite-harzburgite-chromitite complexes as re- (<400 °C). Hence, the Ise serpentinites are considered to fractory residues in the Sangun-Yamaguchi zone, western Japan. Journal of Petrology, 21, 141-165. Arai, S. (1994) Characterization of spinel peridotites by olivine- spinel compositional relationships: review and interpretation. 1.0 Chemical Geology, 113, 191-204. Arai, S. (1997) Control of wall-rock composition on the formation 0.8 of podiform chromitites as a result of magma/peridotite inter- action. Resource Geology, 47, 177-187. Bonatti, E. and Michael, P. (1989) Mantle peridotites from conti- 0.6 nental rifts to ocean basins to subduction zones. Earth and # Planetary Science Letters, 91, 297-311. Hz Dick, H.J.B. and Bullen, T. (1984) Chromian spinel as a petroge- 0.4 20% netic indicator in abyssal and alpine-type peridotites and spa- 15% tially associated lavas. Contributions to Mineralogy and Pe- 0.2 10% trology, 86,54-76. Evans, B. (2004) The serpentinite multisystem revisited: Chryso- FMM tile is metastable. International Geology Review, 46, 479- 0.0 0.0 0.2 0.4 0.6 0.8 1.0 506. Evans, B.W. and Frost, B.R. (1975) Chrome-spinel in progressive TiO, wt% metamorphism: a preliminary analysis. Geochimica et Cos- Figure 6. Cr#-TiO2 model (Pearce et al., 2000) for chromian spinel, mochimica Acta, 39, 959-972. showing the partial-melting degree of the Ise serpentinite's pro- Frost, B.R. and Beard, J.S. (2007) On silica activity and serpenti- toliths. Fields of dunite, harzburgite and calculated melting trend nization. Journal of Petrology, 48, 1351-1368. are after Pearce et al. (2000). Hirauchi, K., Tamura, A., Arai, S., Yamaguchi, H. and Hisada, K. 260 M. Zaki Khedr and S.Arai (2008) Fertile abyssal peridotites within the Franciscan sub- Miyakawa, K. (1982) Low-grade metamorphic rocks of the Hida duction complex, central California: Possible origin as de- marginal belt in the upper Kuzuryu river area, central Japan. tached remnants of oceanic fracture zones located close to a Japanese Association of Mineralogists Petrologists and Eco- slow-spreading ridge. Lithos, 105, 319-328. nomic Geologists, 77, 256-265. Imai, N., Terashima, H., Itoh, S. and Ando, A. (1995) 1994 compi- Muintener, O., Hermann, J. and Trommsdorff, V. (2000) Cooling lation values for GSJ reference samples, “Igneous rock se- history and exhumation of lower-crustal granulite and upper ries". Geochemical Journal, 29, 91-95. mantle (Malenco, eastern central Alps). Journal of Petrology, Ishii, T., Robinson, P.T., Maekawa, H. and Fiske, R. (1992) Petro- 41, 175-200. logical studies of peridotites from diapiric serpentinite sea- Nakamizu, M., Okada, M., Yamazaki, T. and Komatsu, M. (1989) mounts in the Izu-Ogasawara-Mariana forearc. In Proceed- Metamorphic rocks in the Omi-Renge serpentinite melange, ings of the Ocean Drilling Program Scientific Results 125 Hida Marginal Tectonic Belt, Central Japan. The Memoirs of (Fryer, P., Pearce, J.A. and Stokking, L.B. Eds.). Ocean Drill- the Geological Society of Japan, 33, 21-35 (in Japanese with ing Program, College Station, TX, 445-485. English abstract). Ishiwatari, A. (1985) Igneous petrogenesis of the Yakuno ophiolite Niu, Y. (1997) Mantle melting and melt extraction processes be- (Japan) in the context of the diversity of ophiolites. Contribu- neath ocean ridges: evidence from abyssal peridotites. Jour- tions to Mineralogy and Petrology, 89, 155-167. nal of Petrology, 38, 1047-1074. Khedr, M.Z. and Arai, S. (2009) Geochemistry of metasomatized Niu, Y. (2004) Bulk-rock major and trace element compositions of peridotites above subducting slab: a case study of hydrous abyssal peridotites: implications for mantle melting, melt ex- metaperidotites from the Happo-O'ne complex, central Ja- traction and post-melting processes beneath mid-ocean ridg- pan. Journal of Mineralogical and Petrological Sciences, 104, es. Journal of Petrology, 45, 2423-2458. 313-318. Nozaka, T. (2005) Metamorphic history of serpentinite mylonites Khedr, M.Z. and Arai, S. (2010) Hydrous peridotites with Ti-rich from the Happo ultramafic complex, central Japan. Journal of chromian spinel as a low-temperature forearc mantle facies: Metamorphic Geology, 23, 711-723. evidence from the Happo-O'ne metaperidotites (Japan). Con- Okamura, H., Arai, S. and Kim, Y.-U (2006) Petrology of forearc tributions to Mineralogy and Petrology, 159, 137-157. peridotite from the Hahajima Seamount, the Izu-Bonin arc, Khedr, M.Z., Arai, S., Tamura, A. and Morishita, T. (2010) Clino- with special reference to chemical characteristics of chromian pyroxenes in high-P metaperidotites from Happo-O'ne, cen- spinel. Mineralogical Magazine, 70, 15-26. tral Japan: implications for wedge-transversal chemical cha- Parkinson, I.J. and Pearce, J.A. (1998) Peridotites from the Izu- nge of slab-derived fluids. Lithos, 119, 439-456. Bonin-Mariana forearc (ODP Leg 125): evidence for mantle Kodolanyi, J and Pettke, T. (2011) Loss of trace elements from melting and melt-mantle interaction in a supra-subduction serpentinites during fuid-assisted transformation of chryso- zone setting. Journal of Petrology, 39, 1577-1618. tile into antigorite-an example from Guatemala. Chemical Pearce, J.A., Barker, P.F., Edwards, S.J., Parkinson, I.J. and Leat, Geology, 284, 351-362. P.T. (2000) Geochemistry and tectonic significance of perido- Kunugiza, K., Goto, A., Itaya, T. and Yokoyama, K. (2004) Geo- tites from the south Sandwich arc-basin system, South Atlan- logical development of the Hida Gaien belt: constraints from tic. Contributions to Mineralogy and Petrology, 139, 36-53. K-Ar ages of high P/T metamorphic rocks and U-Th-Pb Takazawa, E., Frey, F.A., Shimizu, N. and Obata, M. (2000) EMP ages of granitic rocks affecting contact metamorphism Whole rock compositional variations in an upper mantle pe- of serpentinite. Journal of the Geological Society of Japan, ridotite (Horoman, Hokkaido, Japan): are they consistent 110, 580-590 (in Japanese with English abstract). with a partial melting process? Geochimica et Cosmochimica Kurihara, T. (2003) Stratigraphy and geologic age of the middle Acta, 64, 695-716. Paleozoic strata in the Kuzuryu lake-upper Ise River area of Tsujimori, T., Liou, J.G., Ernst, W.G. and Itaya, T. (2006) Triassic the Hida-gaien terrane, central Japan. Journal of the Geologi- paragonite- and garnet-bearing epidote-amphibolite from the cal Society of Japan, 109, 425-441 (in Japanese with English Hida Mountains, Japan. Gondwana Research, 9, 167-175. abstract). Tsujimori, T. (2004) Origin of serpentinites in the Omi serpentinite Kurokawa, K. (1985) Petrology of the Oeyama ophiolitic complex melange (Hida Mountains, Japan). Journal of Geological So- in the inner zone of Southwest Japan. Science reports of Ni- ciety of Japan, 110, 591-597. gata University Series E (Geology and Mineralogy), 6, 37- Will, T.M., Powell, R. and Holland, T.J.B. (1990) A calculated 113. petrogenetic grid for ultramafic rocks in the system CaO- Machi, S. and Ishiwatari, A. (2010) Ultramafic rocks in the Kotaki FeO-MgO-Al,O3-SiO2-CO2-H2O at low pressures. Contri- area, Hida Marginal Belt, central Japan: peridotites of the butions to Mineralogy and Petrology, 105, 347-358. Oeyama ophiolite and their metamorphism. Journal of the Yokoyama, K. (1985) Utramafic Rocks in the Hida Marginal Geological Society of Japan, 116, 293-308 (in Japanese with Zone. Memoris of the National Science Museum of Tokyo, English abstract). 18, 4-18. Matsumoto, I., Arai, S. and Harada, T. (1995) Hydrous mineral in- clusions in chromian spinel from the Yanomine ultramafic Manuscript received July 6, 2011 complex of the Sangun zone, Southwest Japan. Journal of Manuscript accepted August 8, 2011 Mineralogy, Petrology and Economic Geology, 90, 333-338 Published online October 15, 2011 (in Japanese with English abstract). Manuscript handled by Akira Ishiwatari
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KHEDER~1.txt
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Geochemical Journal, Vol. 25, pp. 335 to 355, 1991
Attainment of solution and gas equilibrium in
geothermal systemsJapanese
HITOSHI CHIBA
Institute for Study of the Earth's Interior, Okayama Univertity,
Misasa, Tottori 682-01, Japan
(Received August 27, 1990; Accepted April 4, 1991)
The geothermal fluids in seven Japanese geothermal systems are tested for attainment of aqueous
and gaseous equilibrium. The pH of fluids in the geothermal reservoir is approximately buffered by the
assemblage K-feldspar-K-mica-quartz. (Na+)/(K+) and (Na+)/ ( activity ratios are ther
modynamically approximated by reactions between albite and K-feldspar, and between albite and anor
thite (or Ca-zeolites), respectively. The (Mg2+)/(K+)2 activity ratio of high temperature geothermal
fluids of Japan can be , represented by the reaction involving Mg-chlorite and K-bearing silicate minerals,
though at lower temperatures other reactions may be responsible. The geothermal fluids are also com
monly saturated with respect to anhydrite and calcite. A small amount of steam loss in the reservoir does
not significantly affect the aqueous composition of the fluids. The partial pressure of CO2 is controlled
by the reaction involving calcite, K-bearing silicate minerals, and albite or Ca-zeolite in geothermal
systems which are not affected by steam loss and dilution. Equilibrium between CH4, C02 and H2 is at
tained at high temperatures but not maintained to lower temperatures in most Japanese geothermal
systems. The H2/H2S ratio is probably equilibrated with Fe-bearing minerals. Gaseous compositions are
very good indicators to identify processes in the geothermal reservoir, such as boiling and dilution. Last
ly, the major aqueous composition and pH of Japanese neutral Na-Cl type geothermal fluid are predic
table if two variables (e.g., temperature and one of the cation activities) are provided.
INTRODUCTION
In the last decade, many geothermal systems
have been assessed and exploited in Japan. A
large number of deep wells have been drilled into
the geothermal reservoirs, and fluids discharged
from them have been analyzed for their
chemical, gaseous and isotopic compositions.
Some of these compositions are now available in
the literature (e.g. Kirishima: Kodama and Naka
jima, 1988; Okuaizu: Nitta et al., 1987). Most
reservoir fluid compositions are Na-Cl dominant
and represent the composition of fluid which in
teracts with reservoir rocks and forms alteration
minerals. Their compositions are likely to be con
trolled by minerals of the rock matrix depending
on the degree of water-rock interaction and also
by the volatiles added to the system from the
magmatic heat source (e.g., Giggenbach, 1984).Studies of reactions between geothermal fluids
and the reservoir mineral assemblage provides
basic information about processes governing
hydrothermal mass transfer in the shallow part
of the earth's crust.
Arnorsson et al. (1983a) examined composi
tions of many Icelandic geothermal waters and
showed that the major element composition is
predictable to as low as 50°C, provided two
parameters, e.g. temperature and chloride con
centration, are given. They concluded that this is
possible due to the attainment of, or close ap
proach to, an overall chemical equilibrium in
geothermal systems. Giggenbach (1980) dis
cussed reactions involving gases in New Zealand
geothermal systems, and concluded that the com
position of fluids reflects close to complete
equilibrium within the system H20, C02, H2S,
NH3, H2, N2 and CH4. The C02 content of the
335
336 H. Chiba
fluid in equilibrium with alteration minerals in
volving calcite and chalcedony is also predicted
by Giggenbach (1981, 1984). In Japanese geother
mal systems, however, the details of reactions
and chemical equilibrium among geothermal
fluids, gases, alteration minerals and reservoir
rocks have not been discussed.
In this study the aqueous speciation and gas
equilibria of neutral Na-Cl type geothermal
fluids from seven Japanese geothermal systems
and some hot spring waters are used to discuss
the controls on the chemical compositions of
geothermal fluid by the reservoir rocks.
DATA SOURCES AND CALCULATIONS
Geothermal well and hot spring data
The locations of geothermal areas and hot
springs examined in this study are shown in Fig.
1. The geothermal wells and hot springs used in
this study are listed in Table 1. The analytical
data and sampling conditions of Japanese
geothermal wells used in this study are given in
the Appendix Table. Mineral assemblages ob
served in drill cores and cuttings have not been
reported in detail for most geothermal systems.
Reported mineral assemblages are briefly sum
marized in Table 2 for some geothermal systems.
The fluid discharged from geothermal wells
are less affected than hot spring waters by boil
ing, mixing with local meteoric water, precipita
tion of minerals or leaching from rock. Thus,
they are better representatives of the fluid
undergoing water-rock interaction at depth.
Complete sets of chemical and physical data, i.e.
chemical compositions of the liquid and steam
phase, discharge enthalpy and sampling condi
tion, are required for the purpose of this study.
Unfortunately, the number of such complete sets
of published data are small, and have been
chosen carefully from the literature. Except for
well It of Okuaizu, wells with high excess en
thalpy discharges are omitted from Table 1 and
subsequent calculations. If the discharge en
thalpy is cited in the literature, it is easy to ex
clude excess enthalpy wells. Even if the discharge
enthalpy is not cited in the literature, a geother140°
130°145'
Nigorikawa
135°
® Hatchobaru0 okuaizu
SNCSYC
Sic• TJC45'
Takigami O0 SC
INC,ISC
STC
irishima 0
30'40°
Kakkonda A
Sumikawa
35°
Fig. 1. Location of geothermal systems considered
in this study. Large circles indicate the location of
geothermal systems, and small dots indicate hot
springs. The symbols used in the following figures are
indicated next to the names of the geothermal
systems.
mal well whose Na/K geothermometer tempera
ture (Fournier, 1979) is far from its silica satura
tion temperature (Arnorsson et al., 1983b) is
excluded because this may indicate the addition
of steam to the discharge within the reservoir.
Geothermal wells in Table 1, except for It of
Okuaizu, are believed to have discharge en
thalpies which agree with the estimated reservoir
temperatures. These fluids can safely be assumed
to exist as a single liquid phase in their reser
voirs. Thus, total discharge compositions (deter
mined from steam fractions) reflect the reservoir
liquid compositions. Okuaizu well It is an exam
ple of an excess enthalpy well and is used only in
gas calculations. The reservoir temperature used
for speciation and gas calculation is estimated by
the chalcedony (< 180°C) (Arnorsson et al.,
1983b), quartz (> 180°C) (Arnorsson et al.,
1983b) or Na/K geothermometer (Fournier,
1979).
Aquifer rock types of geothermal systems in
Table 1 can be roughly grouped into two
categories, volcanic and marine sedimentary
rocks. The host rock can be characterized from
the B/Cl concentration ratio of associated
Solution and gas equilibrium
O V
cu
b0
`nv
0
0 O
0
V
0 CIO
tiG 0
0
U
F.
O
w
G44.1
N
N 3*C4
0 Cd
0
*
x
z*
Uca
U
0
d N
HON
00 00
Cd
0 >4 rA
aN
M
N N
2 2 N_ N N
N N
2 2
00 00
gC." 00 00
vcVi N N
2 2
N Na,, 00w
Q6) Cd
zF.
N
2
00
NM
~I1 N
N
00O~
00
2
2 ~ 2
vi p,
2 '^
kn tn2
N
00
W)
N O N
2 N 2
U x
e.i
z~
z
CO
o
O
zz0
tiZ
CO
o w cnx x
M
2
N00
00
CO
N
O
N Wn
N N " N
2 2
O v~
N
O N
h N
2 2
N N v;
N N
2 2
N M N NO
2
O
N
N 2
00
C' 0\ N
2
00
M Ii
00
2
* *
N
2
00 kn
~ZHx~xHcnv)x~zZrn
4 0.4,
N W N M
E N N N 01 0 00
z W v, x
[t (~ ~, N O QU ~j H E., c0o~MU°M
MocO,=.W UU~~UUU~
t~ ~ N O M 00 to en 00 yj N N yj , , .-. . -~ ..~ -+"U000U,.., UU wh~wwti~x~Z U
Ca*
0
0 x
C30
U y
xxa 0O
m
w
0
v ~
. O p
y 00 1-4
94 V
w V
C ti
atop q~
* * * **
*337
338 H. Chiba
Table 2. Summary of alteration minerals"
Geothermal
SystemQtz Chi Mica Lm/Wa Anh K-feld Cc Py others
Nigorikawa2)
Sumikawa3)
Okuaizu4)
Hatchobaru5)
Takigami6)
Kirishima')+
+
++
+
+
+
++
++
+
+
++
+
+
+
+
++
+
+
++
+
+
+
++
+
+
+sericite
sericite
prehnite
epidote
sericite
prehnite
epidote
Qtz: quartz, Chi: chlorite, Lm/ Wa: laumontitel wairakite, Anh: anhydrite, Cc: calcite, Py: pyrite.
1) The mineral which is explicitly reported to exist at the production level is noted with a "+ " sign.
2) Yoshida (1991)
3) Sakai et al. (1986); Mitsubishi Material Co. (private communication).
4) Nitta et al. (1987).
5) Kyushu Electric Co. (private communication).
6) Hayashi et al. (1988).
7) Kodama and Nakajima (1988).
1000
m100
10
14
0
water0
10 100 1000 ;10000
CI, mg/I
Fig. 2. Relationship between boron and chloride
concentrations of fluids from geothermal wells. Con
centrations are in liquid phase after steam separation.
The lines in the figure stand for molar ratio of B/ Cl.
Data for Nigorikawa and Kakkonda are from Shigeno
and Abe (1987) and Y. Yoshida (private communica
tion), respectively. Symbols are as in Fig. 1.crustal
10He`%-N2magmatic 0 .01 N2
gas (Usu) X
40
+80
+
+,0
+#
+4air
air saturated
.w groundwater
geothermal fluids (Shigeno and Abe, 1983).
Figure 2 is a plot of B against Cl concentrations
in the liquid phase of geothermal well
discharges. Fluids with B / Cl molar ratios be
tween 0.02 and 0.07 (Nigorikawa, Takigami,
Hatchobaru and Okuaizu) discharge from
volcanic reservoir rocks. In Takigami and Hat
chobaru, the geothermal fluids are considered to
be stored in the volcanic rocks (Hayashi et al.,
1988; Manabe and Ejima, 1984). On the other0 20
%-Ar60 80 I.".Ar
Fig. 3. Relative composition of Ar, He and N2 in
geothermal fluids, hot spring waters, mineral spring
waters and volcanic gases. The composition of
magmatic gas is defined by volcanic gases of Mt. Usu,
Hokkaido, Japan (Matsuo et al., 1982). Large plus
signs: volcanic gases of Mt. Usu, small plus signs:
volcanic gases of other volcanoes (Kiyosu, 1985;
Kiyosu and Yoshida, 1988), small dots: gases in
Japanese mineral and hot spring waters (Urabe et al.,
1985), solid squares: fluids in Nigorikawa (Yoshida,
1991), open squares: Okuaizu (Nitta et al., 1987),
solid triangles: Sumikawa (Ueda et al., 1991), open
triangles: Kakkonda geothermal system (Kiyosu and
Yoshida, 1988).
hand, wells in Kirishima, Kakkonda
Sumikawa have B / Cl ratios higher than and
0.07,
Solution and gas equilibrium 339
and are considered to discharge fluids having in
teracted with marine sedimentary rocks. In
Kirishima, the basement rock, Shimanto
Supergroup, is rich in fractures allowing fluid
rock interaction during ascent, though the main
geothermal reservoir is andesitic rocks at
shallower depth (Kodama and Nakajima, 1988).
The aquifer rock types will be discussed later in
relation to the fluid compositions.
The contribution of magmatic gas to the
geothermal systems can be assessed on an Ar
He-N2 diagram (Giggenbach, 1986). Ar and He
are not reactive at hydrothermal conditions. N2
is also not reactive in conditions of Japanese
geothermal systems examined here, since no ap
preciable NH3 is reported. Therefore, these gases
are able to preserve their source signatures after
interaction with reservoir rocks. The gaseous
compositions of Nigorikawa (Yoshida, 1991),
Kakkonda (Kiyosu and Yoshida, 1988),
Sumikawa (Mitsubishi Material Co., private
communication) and Okuaizu (Nitta et al., 1987)
are plotted in Fig. 3 together with those of
Japanese volcanic gases (Kiyosu, 1985; Kiyosu
and Yoshida, 1988) and mineral and hot springs
(Urabe et al., 1985). The magmatic component is
defined in Fig. 3 by the volcanic gas of Mt. Usu
(Matsuo et al., 1982). Figure 3 indicates that the
geothermal gases of Nigorikawa and Okuaizu
have a relatively large contribution of magmatic
gases. In Kakkonda and Sumikawa, geothermal
gases are negligibly affected by a magmatic com
ponent. Some wells in Sumikawa are highly
affected by magmatic component (Ueda et al.,
1991), but they are not used in this study because
of their large excess discharge enthalpy.
Most hot spring waters are affected by boil
ing, mixing of deep fluids with waters of
different origin (e.g. groundwater and steam
heated water), precipitation of minerals from the
water during upflow and/or interaction with the
host rock (e.g., Giggenbach, 1988). Their com
positions do not usually represent the composi
tion of fluid interacting with the rocks in the
reservoir where geothermal fluids are stored. The
chemical composition of over two thousand hot
springs in various Japanese geothermal systemswas compiled by Hirukawa et al. (1977). The hot
springs least affected by shallow processes have
been selected on the basis of the following
criteria: (1) Na/K thermometer temperature
(Fournier, 1979) agrees with chalcedony
(< 180°C) or quartz (> 180°C) saturation tem
perature (Arnorsson et al., 1983b) within 10°C
and (2) the charge balance after aqueous specia
tion calculation is less than 1%. Twenty four of
the over two thousand hot springs cleared the
two criteria and are used in this study. The
speciation calculations were carried out at the
temperatures estimated by the applicable Si02
geothermometer.
For the purpose of comparison, data of
Icelandic and New Zealand geothermal waters
are taken from Arnorsson et al. (1983a) and
Hedenquist (1990), respectively. Hot spring data
in Iceland were selected on the basis of the same
criteria as for Japan. Aqueous speciations were
re-calculated by the code mentioned in the next
section.
Aqueous speciation and gas calculation
The aqueous speciations are calculated using
the code described by Chiba (1990). Forty five
aqueous species are included in the calculation.
It treats only a single liquid phase at tempera
tures from 25 to 300°C. For calculation of a
geothermal well discharge, the steam phase, in
cluding gases, is condensed back to the liquid
phase in proportion to the steam fraction at the
time of sampling. Aqueous speciation of fluid
with excess discharge enthalpy cannot be
calculated by the code used in this study. The
dissociation constants of aqueous species are
adopted from the thermodynamic data base of
SOLVEQ (Reed, 1982). The thermodynamic
data of minerals are from Helgeson et al. (1978).
The calculation of gaseous species follows
the method of Giggenbach (1980). Fugacity
coefficients of gases are assumed to be unity in
the same manner as Giggenbach (1980). C02
mineral equilibrium constants were generated by
SUPCRT (Helgeson et al., 1978) using the 1981
data base. This results in stability relationships
of Ca-Al-bearing minerals that are slightly
340 H. Chiba
different from the original figures in Giggenbach
(1984).
RESULTS OF AQUEOUS SPECIATION AND
GAS CALCULATIONS
Aqueous species
In Figs. 4 and 5, cation/proton activity
ratios are plotted against reservoir temperatures.
Except for the (Mg")/(H')' ratio, other ratios
show simple patterns against reservoir tempera
ture. This suggests that these ratios are controll
x
z
rn
07
6
5
000%
0 • •0
0 tA
_ K-feld~rF
(=rnlq"
0B
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i
0010
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C.r
9
N 8
7 + N
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least squares
fit
0
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et9%) n
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04
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8'2
10
8
6
Fig. 4. The
tion/proton
discharges
recalculate
tion /proton
The dashed
Symbols are as in Fig. 1.4
150 200 250 300
Temperature, °C
temperature dependence of ca
activity ratios of geothermal well
in Japan. The lines in the figures are
d temperature dependences of ca
ratios in Icelandic geothermal waters.
curve in Fig. 4B represents reaction (3).150 200 250 300
Temperature, °C
Fig. 5. The temperature dependence of ca
tion/proton activity ratios of geothermal well
discharges in Japan. The solid curves in the figures are
recalculated temperature dependences of ca
tion/proton ratios in Icelandic geothermal waters.
The lower curve in Fig. SA is a least squares fit of
some of the plotted data (see text). Symbols are as in
Fig. 1.
ed by mineral buffer systems in the reservoir.
As inferred from Fig. 2, the aquifer rock
types of Nigorikawa, Okuaizu, Hatchobaru and
Takigami geothermal systems are volcanic rocks.
Marine sedimentary (or metasedimentary) rocks
are judged to be the aquifer rocks of Kakkonda,
Sumikawa and Kirishima geothermal systems.
Figs. 4 and 5 indicate that the aquifer rock type
does not systematically affect the major element
fluid compositions. The same phenomena were
observed in Icelandic thermal waters (Arnorsson
et al., 1983a) and in water/rock experiments
(Kacandes and Grandstaff, 1989).
Arnorsson et al. (1983a) showed that ca
tion/proton and cation/cation activity ratios of
Icelandic thermal waters follow simple patterns
against reservoir temperatures as low as 50°C.
The patterns are considered to be the results of
silicate mineral buffer systems, which may vary
with temperature but have smooth transitions
Solution and gas equilibrium 341
because of small differences in the A G of
minerals. However, the patterns given by them
cannot be directly compared with the present
results, because some of their thermodynamic
data used for speciation are different from those
used in this study. The curves in Figs. 4 and 5
were obtained by least squares fits of the
.recalculated speciations of Icelandic samples.
Except for (Mg")/(H')', most cation/proton
ratios scatter around the curves obtained from
Icelandic samples. This suggests that the ca
tion/proton ratios in Japanese geothermal
systems may be controlled by silicate mineral
buffer systems similar to Icelandic geothermal
systems. Detailed discussions about silicate
mineral buffer systems follow examination of the
individual geothermal systems, particularly on
effects disturbing the aqueous and gaseous com
position, such as boiling in the reservoir and mix
ing with waters of different origin. The hot
spring waters are not plotted in Figs. 4 and 5 as
they have a large scatter. Though they were very
carefully selected from the literature, their
erratic results indicate that the waters discharged
from Japanese hot springs have been affected by
processes such as boiling, mixing with shallow
water and/or reaction at low temperature with
minerals not accounted for by the curves in Figs.
4 and 5 before they reach the surface. Therefore,
they are not good representatives of fluid interac
ting with rock at reservoir depths.-10
-15
~, -20 Y
° -25
-30
-35 jr lib•v
o.~
000 0
of0
0
P s3 100"P
a
Gaseous species
Gaseous species are treated in the manner de
scribed by Giggenbach (1980). He also con
sidered the effect of steam loss and gain, and the
attainment of equilibrium in the carbon and
sulfur gas systems is discussed following his inter
pretive framework.
Carbon dioxide and methane are carbon
bearing gases whose concentrations are often
analyzed in geothermal studies. Between these
two gases, the following equilibrium often ex
ists:
CO2+4H2=2H2O+CH4. (1)
The analytical equilibrium constants for this100 150 200 250 300
Temperature, °C
Fig. 6. Plot of analytical log K for reaction (1) ver
sus reservoir temperature, illustrating the effect of
steam loss and gain on C02-H2-CH4 gas equilibrium
(Giggenbach, 1980). The arrow shows an example of a
boiling path for a fluid whose gas composition is in
equilibrium for reaction (1) at 245°C. Symbols are as
in Fig. 1.
N
x
01
0
-1
-2
-3A
lot
'Mite
P;-We--O
Am
e-
0 !
~b oo~
100 200 300
Temperature, °C
Fig. 7. Log H2/H2S of total discharge versus reser
voir temperature. The curves are for pyrite-Fe-AI
silicate, pyrite-magnetite and pyrite pyrrhotite buffers
(Giggenbach, 1980). The arrow shows an example of a
boiling path for a fluid whose composition is buffered
by the pyrite-Fe-Al-silicate reaction (2) at 245°C.
Symbols are as in Fig. 1.
reaction (log Kc") of gases from geothermal
wells are plotted in Fig. 6. The curve at the
center corresponds to the equilibrium of reaction
(1) for all species dissolved in a single liquid
phase. The arrow originating from equilibrium
log Kc" at 245°C is an example of a predicted
boiling path, which is calculated assuming
adiabatic continuous vapor loss. Gaseous com
positions of Japanese geothermal systems scatter
widely around the equilibrium curve. Samples
lying in the region between equilibrium and
342 H. Chiba
equilibrium vapor indicate a fluid which gained
excess steam in the reservoir (Giggenbach, 1980).
The fluid of the Okuaizu system has clearly gain
ed excess steam. In contrast, all fluids from the
Kirishima geothermal system plot below the
equilibrium curve, and on a trend indicating
steam loss from a fluid that may have begun boil
ing at about 245'C; the degree of steam lost is in
dicated by the contours of steam fraction (Gig
genbach, 1980). Equilibrium of reaction (1) ap
pears to be attained in the Kakkonda and
Nigorikawa geothermal systems.
Ratios of H2/H2S mole fraction in total
discharge are plotted against reservoir tempera
tures in Fig. 7. Equilibrium H2/H2S ratios for
pyrite-pyrrhotite, pyrite-magnetite and pyrite
Fe-Al-silicate coexistence given by Giggenbach
(1980) are also shown by the curves in Fig. 7.
The arrow originating from the equilibrium
log(H2/H2S) value of pyrite-Fe-Al-silicate at
245'C indicates an example of an adiabatic boil
ing path. Data for Japanese geothermal wells
scatter widely, though some patterns are visible.
The H2/H2S ratios of some geothermal well
discharges must be influenced by steam loss in
the reservoir. The aqueous solubility of H2 is
much lower than that of H2S, because the gas
distribution coefficient of H2 between vapor and
liquid is greater than that of H2S (Giggenbach,
1980). Therefore, steam (and gas) loss from the
reservoir lowers the H2/H2S ratio, and the
downward scatter from the buffer systems shown
in Fig. 7 may be the result of steam loss from the
reservoir fluid. Fluids from the Kirishima
geothermal system again show a simple vapor
loss trend as for the carbon-bearing gases (Fig.
6). Fluid samples which suggest steam loss in
Fig. 6 (higher temperature wells in Takigami and
one well in Sumikawa) also plot below any buffer
ing systems in Fig. 7, supporting steam loss in
these reservoirs. In the Kakkonda system,
H2/H2S ratios of fluids may be controlled by
pyrite-magnetite in some wells and pyrite-Fe-Al
silicate buffers in other wells. The latter buffer
system was empirically deduced by Giggenbach
(1980) according to the following reaction: pyrite + H2+ H20= FeO(silicate) + 2H2S, (2)
where FeO (silicate) means Fe" in Al-bearing
silicate mineral, e.g. chlorite. The H2/H2S ratio
of one well in Sumikawa may also be accounted
for by reaction (2). The pyrite-pyrrhotite buffer
system cannot be totally ruled out for Japanese
geothermal systems since the vapor loss sug
gested for some wells makes a clear determina
tion of the actual buffer system difficult.
CHARACTERISTICS OF INDIVIDUAL
GEOTHERMAL SYSTEMS
Reactions among aqueous species and silicate
minerals appear close to equilibria in Japanese
geothermal systems (Figs. 4 and 5). However,
boiling and steam gain in the geothermal reser
voir affect the gaseous compositions of some
systems and cause the gaseous compositions to
shift from equilibrium states (Figs. 6 and 7).
These results probably reflect the processes occur
ring in the reservoir of individual geothermal
systems. In other words, partial equilibrium is
frequently attained depending on specific condi
tions in each geothermal system. Before discuss
ing the details of reactions controlling the com
positions of geothermal fluids, we must
recognize the processes influencing the fluid com
position of individual geothermal systems. In
this section, based on the degree of partial
equilibrium attained in an individual system, the
characteristics of each geothermal system will be
briefly discussed.
Nigorikawa geothermal system
The C02 content of fluids in Nigorikawa is
much higher than in other systems (Appendix
Table). Limestone is one of the components of
the geothermal reservoir (Sato, 1988). A possible
explanation of the high CO2 flux is attack of
acidic gas of magmatic origin on limestone to
produce CO2. Despite the addition of a large
amount of CO2 gas to the system, the aqueous
species appear to have approached equilibrium
with silicate minerals, except for (Ca2+)/(H+)2
(Figs. 4 and 5). The high flux of CO2 influences
Solution and gas equilibrium 343
the concentration of aqueous Ca species, since a
large amount of CO2 in the liquid phase will
cause calcium in solution to precipitate as
calcite. The deficiency of Ca2+ can be seen in
Fig. 4C. However, activities of other cations and
pH are not significantly affected by the high C02
flux.
The reaction between CH4, C02 and H2 is
very close to equilibrium, indicating that exten
sive boiling in the geothermal reservoir does not
take place. The contribution of magmatic gases
indicated from Fig. 3 in the Nigorikawa system
does not affect the attainment of equilibrium for
reaction (1). Therefore, fluid discharged from
the Nigorikawa system is representative of fluids
stored in the geothermal reservoir, though the
Ca 21 is slightly influenced by the high C02 flux l empirical curves determined by Icelandic ther
mal waters, except for (Mg2+)/(H+)2 of one fluid
sample. The aqueous species are likely controll
ed by silicate mineral assemblages. Gaseous reac
tions are slightly out of equilibrium, indicating a
small steam loss in the fluid from well S-4.
Okuaizu geothermal system
The one sample from Okuaizu is used in this
study as an example of a fluid with an excess
discharge enthalpy. Though the aqueous com
positions are not plotted in Figs. 4 and 5, they
are far from the empirical curves. The aqueous
composition is strongly influenced by the gain of
steam in the geothermal reservoir, as indicated in
Fig. 6. This sample will be omitted from the
discussion of aqueous equilibrium.
Kakkonda geothermal system
The cation/proton ratios of Kakkonda are
very close to the pattern of Icelandic geothermal
waters, except for (Mg2+)/(H+)2 (Figs. 4 and 5).
The (Mg2+)/(H+)2 ratio is controlled by a reac
tion which includes Mg-chlorite, as discussed
later. Thus, the aqueous composition of fluid in
Kakkonda is closely controlled by the silicate
mineral buffers. Most samples of fluid in Kak
konda are in or close to equilibrium with the
gaseous reaction involving CH4, CO2 and H2,
though one well indicates a slight steam gain
(Fig. 6). The H2/H2S ratio of some samples ap
pears to be controlled by the reaction involving
magnetite and pyrite, and others by the reaction
of Fe-Al-silicate and pyrite, as mentioned earlier
(Fig. 7). Unfortunately, these buffer systems can
not be confirmed by field observation since de
tailed study of Fe-bearing minerals has yet to be
reported. Figures 4 to 7 suggest that fluids in the
Kakkonda system are in full equilibrium with the
reservoir rock.
Sumikawa geothermal system
. The Sumikawa geothermal system is located
about 20 km from the Kakkonda geothermal
system. The salinity of fluid (0.006 mole Cl
kg/H20) is the lowest among fluids examined
here. The cation/proton ratios are close to theHatchobaru geothermal system
All aqueous species in the Hatchobaru
system are close to equilibrium. Since the gas
data required for calculating gaseous equilibria
are not available, boiling and/or steam gain are
not assessed in Figs. 6 and 7. However, Pco2 ap
pears to be closely buffered by the mineral
assemblage, as discussed later. This means that
samples used in this study are not likely affected
by boiling or steam gain. The fluid in the Hat
chobaru system, including gases, is in full
equilibrium with the reservoir rocks.
Takigami geothermal system
A plot of Cl concentrations vs. discharge en
thalpies for the Takigami system (Takenaka and
Furuya, 1991) indicates a strong dilution trend
from the parent fluid (TT-14 well) towards
shallow water whose Cl concentration can be
represented by local hot spring waters. Samples
from wells NE-2 and NE-3 (not plotted in
Takenaka and Furuya, 1991) also lie on the same
dilution line originating from the fluid of TT-14;
they are the most diluted well discharges of
Takigami. The aqueous compositions plot close
to the empirical curves with trends crossing the
empirical curves (Figs. 4 and 5). These trends
might be characteristic of fluids diluted by
shallow water (e.g., Hedenquist, 1990), though
344 H. Chiba
the cation/cation activity ratios are close to
equilibria with alteration minerals, as discussed
later. Compositions of carbon-bearing gases
deviate from equilibrium at high temperature
and approach equilibrium at low temperature.
This erratic result may indicate that reaction
among CH4, CO2 and H2 does not take place in
the dilution process at Takigami and that the
composition of carbon-bearing gases in the
parent fluid which had lost steam in the reservoir
is preserved in the diluted fluids. The composi
tion of fluids in the Takigami geothermal system
is affected by progressive dilution across the
system.
Kirishima geothermal system
All cation/proton activity ratios in this
system plot near the empirical curves, but scatter
wider than in other geothermal systems (Figs. 4
and 5). All gas data indicate steam loss in the
reservoir (Figs. 6 and 7), with boiling beginning
at about 245'C. The fluid with the largest steam
loss, as indicated in Fig. 6, shows the largest
deviation from equilibrium conditions in Figs. 4
and 5. This suggests that aqueous compositions
are slightly affected by steam loss in the reser
voir. Among the geothermal systems examined
here, Kirishima appears to be representative of a
system influenced by steam loss.
DISCUSSION
Effects of boiling and dilution on aqueous and
gas compositions
The solute-mineral, gas-gas and gas-mineral
equilibria discussed earlier appear to be most
closely attained in the Kakkonda and Hat
chobaru geothermal systems, whose reservoir
temperatures cover the range of the studied
geothermal systems, except for low temperature
Takigami wells. In Japanese geothermal
systems, gaseous compositions scatter wider
around empirical or theoretical equilibrium
values than aqueous compositions, suggesting
higher sensitivities of gaseous species for boiling
in the reservoir and/or mixing.
Figures 6 and 7 indicate that the fluids in theKirishima geothermal system are influenced by
boiling, as previously mentioned. According to
Fig. 6, the amount of steam lost from the reser
voir is as high as 10% at Kirishima, assuming
that reaction (1) controlled the initial gas com
position. Reed and Spycher (1985) calculated the
boiling effects on a Broadlands-like fluid in a
fully equilibrated system, in which instantaneous
equilibrium is always attained among minerals,
gases and aqueous species. According to them,
ten percent isoenthalpic boiling of an initially
278'C fluid to 243'C causes a pH increase of
about 0.5 unit, but does not significantly change
simple cation activities. Thus, boiling causes the
cation/proton ratios to deviate to higher values
(Figs. 4 and 5). The results in Figs. 4, 5, 6 and 7
support boiling and steam loss in the reservoir,
with the data from Kirishima roughly agreeing
with the temperature-dependent curve if they are
corrected for steam loss and pH increase. Steam
and gas loss clearly affects not only the gaseous
compositions but also aqueous compositions to
a lesser extent.
The effects of dilution by water of shallow
origin can be seen in the Takigami geothermal
system. In considering reactive elements (e.g. Na
and K) versus a mobile element (Cl), fluid com
positions should plot on a mixing line, if there is
no reaction between aqueous species and
minerals after the mixing. However, this is not
the case for the Takigami geothermal system, as
fluid compositions deviate from simple mixing
lines (Takenaka and Furuya, 1991). The devia
tions of fluid compositions from mixing lines sug
gest two possibilities: (i) the mixing is not a sim
ple two endmember mixing or (ii) reactions
between aqueous species and the reservoir rock
take place to some extent after mixing. In the
Takigami geothermal system, reservoir tempera
tures calculated by the Na-K-Ca geother
mometer agree well with those of the Si02
geothermometers (Appendix Table), and Ca"
and SO4 concentrations of fluids are saturated
with respect to anhydrite (Hayashi et al., 1988).
Therefore, the second possibility is favored for
the Takigami geothermal system, though
equilibria between aqueous species and rock are
Solution and gas equilibrium 345
a.
a.9
8
7
6
5
4
9
8
7
6
5JapanA
Q-01 O'o~ 0.01
pO
chalcedonyquartz
lvt~ba ~9--; 000,
Qr ..
Iceland
chalcedonya•
quartzB
0.01
9.1
0 100 200
Temperature, °C
dependence300
Temperature
H20.Fig. 8. A: of pH in
Japanese thermal waters. Lines indicate the tempera
ture dependence of pH when pH is buffered by the K
feldspar-K-mica-quartz (or chalcedony at less than
200°C) assemblage at a Na+K concentration of 0.1
and 0.01 mole/kg Symbols are as in Fig. 1.
B: Temperature dependence of pH of Icelandic ther
mal waters. Large circles indicate well discharges.
Small dots represent hot spring waters.
not complete, as indicated by the scatter of ca
tion/proton activity ratios (Figs. 4 and 5). The
gas composition also appears to be out of
equilibrium (Figs. 6 and 7).
Control of pH of Japanese thermal waters
mole/kg H20
K-feldspar + H + The pH values at reservoir temperatures are
calculated for Japanese thermal waters and are
plotted in Fig. 8A. The pH of Icelandic thermal
waters are also plotted in Fig. 8B for com
parison. The lines in Fig. 8 indicate a tempera
ture dependence of pH, assuming (i) concentra
tion of E (Na + K) indicated in
Fig. 8, (ii) the relation between activity ratio of
Na+ / K+ and temperature applies (Arnorsson et
al., 1983a) and (iii) the following reaction:
1.5
= 0.5 K-mica + 3 quartz(or chalcedony) + K+ .
(3)Chalcedony saturation was assumed at less than
200°C. The assumed Na+/K+ activity ratio is
very close to a fluid which is in equilibrium with
low albite and microcline (Arnorsson et al.,
1983a).
(Na + K)
H20,
mole/ H20
H20
H20, Geothermal wells in Japan have pH values
around the line of concentration of
0.01 mole/kg except for wells in
Nigorikawa. As the E (Na + K) concentrations
of fluids from these wells range from 0.009 to
0.05, the reservoir pH values appear to be con
trolled by reaction (3). The reservoir pH of
Nigorikawa is lowest among all wells in this
study. Wells in Nigorikawa discharge fluid with
the highest concentration of E (Na + K) and E
CO2 in Japan. The E CO2 concentration is as
high as 0.14 kg and E (Na + K) con
centration is 0.21 mole/kg on a total
discharge basis, with the high CO2 probably
related to the limestone-bearing basement rocks.
The reservoir pH of Nigorikawa is close to the
pH predicted by the silicate mineral buffer
system for the measured E (Na + K) concentra
tions of 0.21 mole/kg indicating the high
E CO2 concentration does not affect the reser
voir pH. Therefore, the reservoir pH of fluids
discharged from Japanese geothermal wells are
approximately buffered by a K-bearing silicate
mineral assemblage.
C02 to
(Na + K)
Na + K The pH values of Japanese hot spring waters,
especially those of high temperature, show a
much larger scatter around the pH mineral
buffer than Icelandic thermal waters. This may
result from the loss of a small amount of
the vapor during boiling to the surface. On the
other hand, the pH values of Icelandic hot
spring waters follow the theoretical temperature
concentration. dependence at low E
This is expected because most Icelandic hot
concenspring waters have uniformly low
trations of around 0.001 mole/kg H2O. Also,
Icelandic well discharges have pH values close to
the pH predicted for their E (Na + K) concentra
tion, assuming a K-bearing silicate mineral
buffer.
346 H. Chiba
Control of aqueous species by silicate mineral
buffers
Saturation of silica minerals, such as quartz
or chalcedony, is implicitly assumed in this
study, and is used to estimate the reservoir tem
peratures of geothermal fluids. The relatively
fast reaction rate between silica minerals and
fluid at geothermal temperature is widely ac
cepted (e.g., Ellis and Mahon, 1977). Also, there
is little silica precipitated during the rapid ascent
of fluid from the bottom of the well to the sur
face. Furthermore, the silica temperatures of
most fluids used in this study agree well with
Na/K temperatures (Appendix Table), suppor
ting the assumption of silica saturation at reser
voir temperatures.
The saturation state of individual Al-bearing
silicate minerals is not easily determined from
the component aqueous species because the ac
curacies of Al analyses in geothermal fluids and
the dissociation constants of Al(OH)n3-n>+
species are frequently not adequate for such
evaluation (Chiba, 1990). Some examples of in
dividual mineral saturations are discussed for
geothermal well discharges by Chiba (1990). Fur
ther investigations about the saturation state of
individual Al-bearing silicate minerals are not
considered in this paper, since the quality of Al
analyses for the geothermal fluid samples studied
here is uneven.
However, we can investigate the reactions
controlling the relative abundance of aqueous
species using cation/cation activity ratios. In
Fig. 9, cation/cation activity ratios of Japanese
geothermal fluids are plotted against reservoir
temperatures. Both (Na+) / (K+) and
(Na+) / (Ca2+) show simple patterns against
reservoir temperature and agree with the em
pirical curve obtained from Icelandic thermal
waters.
The dashed curve in Fig. 9A, calculated from
the thermodynamic data of Helgeson et al.
(1978), is the temperature dependence of the
(Na+)/(K+) ratio of the following reaction:
albite + K+ = K-feldspar + Na+. (4)
The close agreement of Japanese geothermal2.0
Y 1.5
z
0 1.0
0.5
c 1
z
0 0
-1
1 N
Y
0 N
0)
~ -1
s
-2
-3+
150A
®r~
rr~lo
I
B
i a.---------- i'
Anorsson et al. (1983)
I
Jb, Ac ~ele ~tC
O
O
9L*: i9i
Aot0
41e
Fig. 9. The temperature f
tion activity ratios
Solid curves marke
recalculated tempe f
tion/proton ratios 'The dashed curve i p
dependence of (Na+)/(K ff y o
( ). )
Dashed curve in Fig. 9B corresponds to combinations
of reactions (5) at <250 and (6) at > 250°C. The other
two curves correspond to reaction (8); the upper at
chalcedony and the lower at quartz saturation. Sym
bols are as in Fig. 1.200 250 300
Temperature, 00
dependence o cation/ca
of Japanese geothermal fluids.
d Arnorsson et al. (1983) are
rature dependences o ca
of Icelandic geothermal waters.
n Fig. 9A indicates tem erature
+ bu ered b reacti n 4
fluids with the dashed curve indicates the
(Na+) / (K+) ratio can be represented by the reac
tion of low albite and K-feldspar, as pointed out
in other geothermal systems (e.g. Arnorsson et
al., 1983a; Giggenbach, 1988).
The (Na+ / (Ca2+) ratios (Fig. 9B) show
strong correlations with temperature, except for
the Nigorikawa geothermal system, and also
Solution and gas equilibrium 347
agree well with the empirical curve obtained by
Icelandic thermal waters. The scatter of
(Na+)/ (Ca2+) ratios around the curve is much
smaller than that for Icelandic thermal waters.
The Nigorikawa (Na+) / (Ca2+) ratios are
distinctly high compared to those of other
geothermal systems. As (Na+)/ (H+) ratios of
Nigorikawa have the same pattern as those of
other geothermal systems, the deviation of
(Na')/1(_Ca') is due to low (Ca")/(H')'. The
low (Ca2+)/(H+)2 is caused by the high C02 flux
in Nigorikawa system, as discussed previously.
The dashed curve in Fig. 9B is a combination of
the following two reactions:
2 albite + Ca 2+ +4H20 =
laumontite + 2Na+ + 2 quartz (<2501Q, (5)
and
2 albite + Ca' + 2H20
wairakite + 2Na+ + 2 quartz (>250'C). (6)
The dotted curve corresponds to the reaction:
2 albite + Ca2+ = anorthite + 2Na + + 4 quartz.
(7)
Above 200°C, the (Na+)/ (Ca2+) ratios of
Japanese geothermal fluids plot close to the dot
ted curve, indicating the (Na+)/ (Ca2+) ratio of
geothermal fluids can be represented by reaction
(7). Reaction (6) also results in similar
(Na+)/ (Ca +) ratios to reaction (7) above
250°C.
Other information about aqueous species
relating to equilibrium with alteration minerals
can be obtained from the Na-K-Mg triangular
diagram (Giggenbach, 1988). The chemical com
positions of geothermal wells are plotted in this
diagram (Fig. 10), except for Kirishima wells,
where Mg concentrations were not reported.
Also plotted as small dots are Icelandic thermal
waters used to define the curves on Figs. 5, 6 and
9. According to Giggenbach (1988), when full
equilibrium is attained between geothermal fluid
and alteration minerals, reaction (4) and the
following reactions control Na, K and Mg con
centrations:K/100%-Na 60
40
rk:Na/1000
full equilibrium
a ° $° • o
Q • °
immature waters
iMg
Fig. 10. p g
trations of Japanese geothermal fluids and Icelandic
thermal waters. Small dots indicate Icelandic thermal
waters. Large symbols are as in Fig. 1. The region of
immature waters corresponds to the composition of
waters related to rock dissolution without any control
by mineral equilibria (Giggenbach, 1988).0 20 40 60 80 100
-Mg
Relationshi between Na, K and M concen
0.8 K-mica + 0.2 Mg-chlorite + 5.4 silica + 2 K+
=2 .8 K-feldspar + 1.6 water + Mg2+ . (8)
As indicated from Fig. 10, geothermal fluids in
five Japanese geothermal systems (Nigorikawa,
Kakkonda, Sumikawa, Okuaizu and Hat
chobaru) are at or very close to full equilibrium.
However, geothermal fluids of Takigami and
one sample from Sumikawa are in partial
equilibrium in terms of combination of reactions
(4) and (8). Also, about half of Icelandic thermal
waters are far from full equilibrium. Figs. 5A,
9C and 10 suggest that reaction (8) i-s not attain
ed in fluids of Takigami, Sumikawa and some
Icelandic geothermal systems.
In Fig. 9C, two curves are shown in addition
to the empirical curve obtained by. Icelandic ther
mal waters. These curves correspond to reaction
(8); the upper curve uses chalcedony (Giggen
bach, 1988) and the lower one uses quartz as the
silica phase. The fluids, which are indicated to be
in, or close to, full equilibrium (Fig. 10), plot
close to the lower curve, suggesting that their
(Mg2+)/(K+)2 ratios are controlled by reaction
(8) in the presence of quartz. In Fig. 5A an alter
native curve can be drawn by a least squares fit,
connecting fluids in full equilibrium (Fig. 10)
(i.e., all wells of Hatchobaru, Nigorikawa and
348 H. Chiba
-4
-6
-8
a
o -10
-12
-14
-16
-6
-8
a
rn 10 0
-12
-14
-16
-6
-8
a
rn 10 0
-12
-14
-16
0 100 200 300
Temperature, °C
Fig. 11. Plot of logarithm of activity product (log
Q) versus reservoir temperature. Solid circles stand
for activity products of Ca" and SO,', in geothermal
fluids, and open circles for Ca" and COj-. Small and
large circles indicate hot spring and geothermal well
samples, respectively. The upper line is the solubility
product of anhydrite and the lower is calcite. From
top to the bottom, figures are for Japanese, Icelandic
and Broadlands, N. Z., thermal waters, respectively.A supersaturated
r
undersaturated 0
Japan
II I
Bsupersaturated
hte
o 0
undersaturated
Iceland
C supersaturated
N,
0 e0 0 ,a
0
undersaturated
Broadlands, N.Z0C
Kakkonda, and one in Sumikawa). The new
curve (Fig. 5A) may represent the (Mg2+)/(H+)2
ratio when the fluid is buffered by reaction (8) .
The Icelandic empirical curve may correspond to
another reaction applicable at lower tempera
tures, since many Icelandic thermal waters at
temperatures lower than 200°C plot very close to
this curve.
In Fig. 4B the dashed line indicates the tem
perature dependence of the (K+)/ (H+) ratio of
reaction (3), which was calculated from thermodynamic data (Helgeson et al., 1978), assum
ing that activities of all solid phases are unity.
The close agreement of the calculated (K+)/(H+)
ratios of geothermal fluids to the dashed line at
high temperature suggests that the relative abun
dances of K+ and H+ ions are controlled by reac
tion (3), as expected from the role it plays as a
pH buffer of geothermal fluids.
Saturations with respect to anhydrite and calcite
Activity products, (Ca2+)(SO4-) and (Ca2+)
(C02 3 ), are plotted (as log Q) against reservoir
temperature in Fig. 11. The upper line of each
figure indicates the solubility product of
anhydrite and the lower is calcite. All Japanese
geothermal wells are approximately saturated
with respect to anhydrite (Fig. 1 1A). Hot spring
waters are slightly undersaturated, but are much
closer to saturation than Icelandic hot springs
and low temperature geothermal fluids. Geother
mal fluids in Broadlands, New Zealand, are ob
viously undersaturated with respect to
anhydrite, even at high temperature.
Sakai and Matsubaya (1974) categorized
Japanese thermal waters into four types based
on their stable isotopic ratios. One is Green-tuff
type thermal water, which leaches fossil Miocene
marine sulfates that precipitated contem
poraneously with the Green tuff formation.
Most of the Japanese hot springs studied here
are located near the Green tuff formation. In
these cases, approach to anhydrite saturation of
Japanese hot spring waters may be due to the
leaching of fossil marine anhydrite. In some
geothermal systems, however, there is no indica
tion of the influence of fossil seawater on the
composition of geothermal fluid. For example,
in Takigami and Hatchobaru, the B / Cl ratios
(Fig. 2) suggest that the geothermal reservoirs
consist of volcanic rocks. Anhydrite is abundant
in veins of altered rocks at Takigami, and the
aqueous sulfate concentration is controlled by
solubility of anhydrite (Hayashi et al., 1988).
Anhydrite is also found in veins in all of the
other geothermal systems studied here (Manabe
and Ejima, 1984; Yoshida, 1990; Mitsubishi
Material Co., private communication; Kodama
Solution and gas equilibrium 349
and Nakajima, 1988). Therefore, the sulfate con
centrations of geothermal fluids must be controll
ed by anhydrite in the reservoir, similar to that in
Takigami. The common attainment of anhydrite
saturation is a significant characteristic of
Japanese geothermal systems when compared to
systems in Iceland and Broadlands, N.Z. The re
maining question is why anhydrite commonly
forms in Japanese geothermal systems. A stable
isotopic study of anhydrite and dissolved sulfate
may provide the clues to understanding the
widespread presence of anhydrite in Japanese
geothermal systems.
All the fluids discharged from Japanese
geothermal wells in this study are approximately 'saturated with respect to calcite, as expected
from its common presence in drilling cores and
cuttings of geothermal wells. About half the
Japanese hot spring waters are super-saturated
with respect to calcite, whereas Icelandic hot spr
ing waters are just saturated. The surface temper
atures of hot springs whose waters are super
saturated with respect to calcite are close to boil
ing. If a small amount of vapor is lost by surface
or subsurface boiling, much C02 escapes from
the liquid to steam phase because of the large gas
distribution coefficient (Giggenbach, 1980). This
removal of CO2 from the hot spring before it is
sampled causes an increase in pH and C03-,
which results in an increase in the calculated
(Ca2+)(CO3-) activity product. The calculation in
dicates qualitatively that the observed super
saturation of hot spring waters with respect to
calcite results from C02 loss, probably due to
boiling.
Anhydrite and calcite might control gas
equilibrium through the following reaction:
calcite + H2S + H20(1) = anhydrite + CH4. (9)
The logarithm of equilibrium constants, i.e. log
PCH4/PH2s are almost temperature independent at
around -4.9 (Giggenbach, 1980). Log
PCH4/Px2s values of Japanese geothermal well
discharges range from -0.9 to 1.9, indicating
that equilibria are attained among Ca2+, C03
and SO4 ions only in the aqueous phase, and
that gases are far from equilibrium involvinga9
CD 0-2
-1
0
1
2
3 prehnite
i
0100 grossuiar
v
Cap
pyrophyllite
Fig. 12. c
voir temp
minerals f gg
The solid curve marked (C) indicates the temperature
dependence of Pco1 buffered by reactions (9) and (10).
The curve marked (Q) is for reactions (10) and (11) or
(12). Symbols are as in Fig. 1.100 150 200 250 300
Temperature, °C
o2 (bars) and reserRelationship between P
erature. Stability fields of Ca-bearing silicate
recalculated a ter Gi enbach (1980). are
anhydrite and calcite. The equilibrium of reac
tion (9) is also not attained in the geothermal
systems studied by Giggenbach (1980).
Therefore, the reaction rate among anhydrite,
calcite and carbon and sulfur-bearing gases
must be very sluggish.
Partial pressure of CO2
Partial pressures of CO2 are plotted against
reservoir temperatures in Fig. 12. The stability
relations among Ca-bearing minerals are from
Giggenbach (1984), and are drawn assuming the
presence of calcite and chalcedony (i.e. about
10% super-saturated with respect to quartz).
Two solid curves calculated from ther
modynamic data (Helgeson et al., 1978) repre
sent the potential C02 buffer systems involving
K-feldspar and K-layer silicate in a fully
equilibrated system. The lower solid curve mark
ed (C) is a combination of the following two reac
tions:
laumontite + K-feldspar +C02=
calcite + K-mica + 4 chalcedony
+ 3H20 (<250°C), and (10)
wairakite + K-feldspar + CO2 =
calcite + K-mica + 4 chalcedony
+H2O (>250°C). (11)
350 H. Chiba
The upper curve (Q) assumes the presence of
quartz instead of chalcedony. The corre
sponding reactions are as follows:
laumontite + K-feldspar +C02=
calcite + K-mica + 4 quartz
+3H20(< 250'C), and (12)
wairakite + K-feldspar + CO2 =
calcite + K-mica + 4 quartz
+H2O (>250°C). (13)
The following reaction also gives similar Pco2
values to reaction (13) between 250 and 300°C:
anorthite + K-feldspar + C02+ H20=
calcite + K-mica + 2 quartz. (14)
The geothermal fluid of Okuaizu, which has ex
cess enthalpy discharges, has a very high Pco2, as expected from the gain of excess steam. Wells in
Nigorikawa discharge fluids of high C02 con
tent. Fluids from Kirishima, which show a
strong boiling trend in Figs. 6 and 7, are also
best explained by C02 loss accompanying steam
loss. The near horizontal distribution of Pco2 in the Takigami system (Fig. 12) can be accounted
for by dilution of high temperature fluid by a
cooler water, with boiling and gas loss thus being
quenched.
Only the fluids from Hatchobaru plot on the
solid curve (Q). Since temperatures calculated by
the quartz geothermometer for fluids from the
Hatchobaru system agree with Na/K tempera
tures, the prevailing silica phase is expected to be
quartz. Thus, the Pco2 of Hatchobaru wells is
probably controlled by the CO2 buffer system of
reactions (13) or (14). All wells in Kakkonda and
a low temperature well in Sumikawa plot be
tween curves (C) and (Q). Zeolitic minerals have
errors in the free energies (e.g., Hedenquist and
Browne, 1989), such that the positions of curves
(Q) and (C) cannot be well-defined at present.
The Pco2 of wells in the Kakkonda and
Sumikawa systems may be buffered by reaction
(10) or (12). The Pco2 of Japanese geothermal
fluids which are not greatly affected by boiling or
mixing are probably controlled by silicatemineral buffers described by reactions
through (14).(10)
Chemical condition of the neutral pH geother
mal reservoir
The pH of the neutral pH, NaCl-dominant
geothermal fluids can be approximated by reac
tion (3) (Fig. 8). Activity ratios of major cations,
such as (Na+) / (K+), (Na+) / (Ca2+) and
(K+)2/(Mg2+), are likely controlled by a silicate
mineral assemblage above 150°C or more conser
vatively between 200° and 300°C (Fig. 9).
Japanese geothermal fluids are also saturated
with respect to anhydrite and calcite (Fig. 11). At
tainment of these equilibria suggest that the
chemical compositions of geothermal fluids are
roughly predictable if two parameters are provid
ed, as stated by Arnorsson et al. (1983a). For ex
ample, if temperature and Z (Na + K) concentra
tion are provided, pH and activities of Na+, K+,
Ca2+, Mgt+, SO4 and C02 3 ions can be
calculated directly from Figs. 8, 9 and 11, since
Na+ and K+ ions are the most abundant among
the Na and K-bearing species. Si02 (aq) concen
tration can be fixed by the saturation with
respect to quartz or chalcedony. The concentra
tion of the most abundant anion, Cl-, can be
also calculated from a charge balance equation
after iterative calculations involving complex
aqueous species.
The oxygen fugacity range of fluid in
equilibrium with volcanic rocks between 200°
and 500°C were experimentally determined by
Kishima (1989) as a function of temperature.
The oxygen fugacities of neutral pH geothermal
fluids almost agree with the experimental data
(Chiba, unpublished data). This means that the
oxygen fugacity as well as the full chemical condi
tion in the neutral pH geothermal reservoir can
be roughly predicted if two parameters are pro
vided. The chemical condition predicted in this
way can be used as a constraint for chemical
modeling of geothermal systems, as a chemical
condition for simulation calculation of transport
of a particular element, and so on.
Solution and gas equilibrium 351
CONCLUSIONS
The aqueous species and pH of fluid discharg
ed from Japanese geothermal systems are approx
imately in equilibrium with their reservoir
minerals. The reservoir rock type as well as the
salinity of fluid does not affect the equilibrium
composition of aqueous species. Boiling with as
much as 10% steam loss in the geothermal reser
voir only slightly affects the activity ratios of
aqueous species. The pH and (K+)/(H+) activity
ratio of fluid in the geothermal reservoir is likely
controlled by K-bearing silicate minerals.
(Na+) / (K+) and (Na+) / ( activity ratios
are thermodynamically approximated by reac
tions between albite and K-feldspar, and be
tween albite and anorthite (or Ca-zeolites), re
spectively. The (Mg2+)/(K+)2 activity ratio of
Japanese high temperature geothermal fluid can
be represented by reaction among Mg-chlorite,
K-bearing silicate minerals and quartz, though at
lower temperatures other reactions may be
responsible for controlling this ratio. In addition
to silicate minerals, Japanese geothermal fluids
are saturated with respect to anhydrite and
calcite.
The partial pressure of CO2 is controlled by
reactions involving calcite, K-bearing silicate
minerals, and albite or Ca-zeolite in geothermal
systems not affected by steam loss and dilution.
Equilibrium between CH4, C02 and H2 is attain
ed at high temperatures but not maintained to
lower temperatures in most Japanese geothermal
systems, due to steam and gas loss, and
sometimes dilution. The H2/H2S ratios of some
fluids are likely equilibrated with Fe-bearing
minerals, though detailed studies of Fe-bearing
minerals are required to confirm the buffer
system of sulfur-bearing gases. Gas composi
tions are very good indicators of processes in the
geothermal reservoir, such as boiling and dilu
tion.
The characteristics of individual geothermal
systems can be described by the degree of partial
equilibrium attained in each system. Such
descriptions can become more quantitative if
kinetic data of reactions can be incorporated. Finally, the major aqueous composition and
pH of Na-Cl type geothermal fluids in Japan are
predictable if two variables (e.g. temperature
and one of the activities of the major com
ponents) are provided. Also, the results of this
study provide the basic knowledge necessary to
investigate the processes governing hydrother
mal mass transfer in the shallow part of the
earth's crust.
Acknowledgments-I wish to thank W. F. Giggen
bach, J. W. Hedenquist, N. Takeno and H. Shigeno
for their constructive comments on an earlier version
of this manuscript. I also thank K. Shimada, T.
Yokoyama, A. Ueda, T. Takenaka, Y. Yoshida,
Kyushu Electric Co., Mitsubishi Material Co. and
Idemitsu Kosan Co. for supplying much of the un
published data that made this study possible.
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Solution and gas equilibrium
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Chiba (1990) Attainment of solution and gas equilibrium in japanese geothermal systems.txt
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Chemical Geology, 68 (1988) 317-327 317
Elsevier Science Publishers B.V., Amsterdam -- Printed in The Netherlands
[5]
ASTHENOSPHERIC INJECTION AND BACK-ARC OPENING"
ISOTOPIC EVIDENCE FROM NORTHEAST JAPAN
SUSUMU NOHDA 1, YOSHIYUKI TATSUMI 2, YO-ICHIRO OTOFUJI :~,
TAKAAKI MATSUDA 4 and KYOICHI ISHIZAKA 5
IKyoto Sangyo University, Kamigamo, Kita-ku, Kyoto 603 (Japan)
2Department o[ Geology and Mineralogy, Faculty of Science, Kyoto University, Kyoto 606 (Japan)
:~ Department of Earth Sciences, Kobe University, Kobe 657 (Japan)
4Department of Geology, Himeji Institute of Technology, Himeji 671-22 (Japan)
5Institute of Earth Science, College of Liberal Arts, Kyoto University, Kyoto 606 (Japan)
(Received August 8, 1987; revised and accepted January 6, 1988)
Abstract
Nohda, S., Tatsumi, Y., Otofuji, Y., Matsuda, T. and Ishizaka, K., 1988. Asthenospheric injection and back-arc open-
ing: isotopic evidence from northeast Japan. Chem. Geol., 68: 317-327.
Nd and Sr isotopic compositions were determined for the Tertiary volcanics from the back-arc side of the NE Japan
arc. Sr isotopes show a linear trend through time from an enriched signature (STSr/S6Sr=0.705437) to a depleted
signature ( STSr/SGSr = 0.70270). In a complementary fashion, Nd isotopes start at low value ( eNd = -- 0.80) and show
a gradual increase (~Nd = 8.3 ) with decreasing age. Isotopic change of Nd and Sr from the enriched signature to the
depleted one is synchronous with the opening of the Japan Sea at ~ 15 Ma. This synchronism indicates that the
opening of the Japan Sea was initiated by the injection of the asthenosphere. During the pre-opening stage, the mantle
wedge was composed of a two-layered structure: the sub-continental lithosphere and the underlying asthenosphere.
The volcanics of this stage characterized by an enriched isotopic signature were derived from a source with a higher
proportion of sub-continental lithosphere. The sub-continental lithosphere of the back-arc side was thinned by the
injection of the depleted asthenosphere, which accelerated the growth of the MORB source within the mantle wedge
of the back-arc side and resulted in magma generation with the depleted isotopic signature of Nd and Sr at the post-
opening stage.
1. Introduction
The NE Japan arc and its back-arc basin (the
Japan Sea) consist of one geotectonic unit in
the framework of plate tectonics (Fig. 1 ). The
Japan Sea basin is considered to have formed
by back-arc spreading (Isezaki, 1986). Recent
paleomagnetic work indicates that NE Japan
has rotated counter-clockwise through 47 ° be-
tween 21 and 14 or 11 Ma ( Otofuji et al., 1985a) and SW Japan was subjected to a clockwise ro-
tation of 60 ° at ~ 15 Ma (Otofuji and Matsuda,
1983 ). These differential rotations are inferred
to have caused the opening and formation of
the Japan Sea basin (Otofuji et al., 1985b). It
suggests that NE Japan had been a part of the
eastern margin of the Asian continent before
the rotation.
The opening of the Japan Sea implies a
change of the Japan region from continental
0009-2541/88/$03.50 © 1988 Elsevier Science Publishers B.V.
318
Fig. 1. Index map of NE Japan. Localities of sampling sites
are represented by the star marked AKITA. Small solid cir-
cles indicate Quaternary volcanoes. The larger solid circle
denotes the Kampuzan volcano. NA=North American
plate; EUR=Eurasian plate; PC=Pacific plate; PHS=
Philippine Sea plate.
margin environment to the present island arc
setting. A large amount of mass movement must
have taken place in the upper mantle of the
spreading region during the back-arc opening
(Garfunkel et al., 1986), and this movement
would have changed the structure of the mantle
and crust beneath the Japan region. Such
structural and compositional changes should be
seen in the isotopic and geochemical character-
istics of the volcanic rocks generated during the
event. Nohda and Wasserburg (1986) reported
the presence of age-dependent variations in the
Nd and Sr isotopic compositions of the Tertiary
volcanic rocks from the back-arc side of the NE
Japan arc, and attributed it to the process of
crustal thinning accompanying the back-arc
opening of the Japan Sea.
To confirm the temporal variation of Nd and
Sr isotopes found by Nohda and Wasserburg
(1986), we made further analyses on the sam-
ples collected from the same area. Based on the
results, we present injection of the astheno-
sphere into the mantle wedge and the subse-
quent thinning process of the sub-continental
lithosphere in order to explain the temporal variation of Nd and Sr isotopes of the Tertiary
volcanics from the back-arc side of the NE Ja-
pan arc.
2. Geologic background
The concept of back-arc opening is estab-
lished in intra-oceanic arcs such as the Mar-
iana and Scotia arcs (Kobayashi and Nakada,
1987; Barker and Hill, 1981; Hussong and
Uyeda, 1981; Chamot-Rooke et al., 1987), where
the spreading initiated just beneath the vol-
canic arc. In contrast, the Japan Sea basin
formed as a result of rifting and lateral move-
ment of the continental fragment. The rifting
had initiated just behind the volcanic arc. As a
result, successive volcanic records are pre-
served in the coast region of the Japan Sea in
NE Japan. In relation with tectonic transition
of the NE Japan arc, we divide the volcanic ac-
tivity of the region into three stages: pre-open-
ing ( > 16 Ma), syn-opening (16-14 Ma) and
post~opening ( < 14 Ma).
The area on the southeast of Akita (30 X 60
km ) is mapped in detail and offers the best field
for the present purpose (Ozawa et al., 1979a,
b). Volcanic activity of the region is summa-
rized as follows. The pre-opening stage is char-
acterized by having a larger volume of rhyolitic
welded tuffthan the contemporaneous andesite
with a minor amount of basalt ( Sugimura et al.,
1963). There are scarcely any reliable age data
for the volcanic activity of the syn-opening
stage. In the post-opening stage, andesitic
products were abundant with minor amounts of
basalt and welded tuff. Further investigation is
required for the mutual genetic relationship
among the various rock types in each stage. We
have collected samples from almost all the ba-
saltic rocks that appear fresh and suitable for
the present purpose, and it seems that no more
useful data points can be obtained from this
region.
3. Analytical procedure
Rock samples of 5-10 kg were roughly
crushed. Fresh fragments were selected, rinsed
TABLE I
Major-element compositions and Nd, Sr isotopic compositions of the Tertiary volcanics from the Japan Sea side of the NE Japan arc
Sample No.
Formation
Locality AKT 12.2 AKT 2.3 NS-833 AKT. 7
Ukibuta basalt in Daisenyama Kanotsume Taikura basalt in
Sugota Fm. Hatamura Fm.
lat.( .... N) 391901 390237 391926 391014
long.( .... E) 1402120 1402138 1401153 1401742
Rock type olivine basalt olivine basalt olivine basalt apbyric basalt NS-839 HS-27
Sanzugawa
39 09 52 39 00 02
140 17 20 140 28 49
olivine basalt hypersthene augite
andesite
Age (Ma) .1 22 (20) 20.9 17.6
SiO2 50.24 50.74 52.04 48.88
TiO2 0.98 2.07 1.69 1.85
Al20:~ 16.42 16.59 16.72 16.54
FezO3 .2 8.91 10.23 8.98 10.47
MnO 0.14 0.17 0.12 0.14
MgO 8.44 4.72 4.84 5.51
CaO 9.22 8.45 7.36 8.50
NaeO 2.66 3.10 3.48 3.31
K20 0.54 1.43 1.36 0.77
P20~ 0.19 0.64 0.48 0.49
2: 97.74 98.14 97.07 96.46
Sr (ppm) 285 456 488 491
Rb (ppm) 7 27 28 9
SVSr/~Srp*:~ 0.705149 _+ 0.000020 0.705476 _+ 0.000019 0.705099 ± 0.000015 0.704674 + 0.000021
STSr/SGSr~ 0.705127 0.705428 0.705051 0.704661
~4aNd/144Nd*a 0.512676±0.000029 0.512586+0.000017 0.512626+0.000018 0.512718±0.000016
eNd 1.05 -- 0.70 0.08 1.87
~s, 8.90 13.17 7.82 2.29 (10) 4.7
48.15 60.05
1.11 0.87
16.72 17.30
8.80 7.26
0.16 0.24
7.49 1.79
9.64 6.28
2.67 4.34
0.71 1.28
0.24 0.26
95.69 99.67
308 334
11 28
0.703452 ± 0.000016 0.703840 __+ 0.000026
0.703438 0.703824
0.512964__+0.000029 0.512855__+0.000027
6.67 4.55
- 15.07 -9.60
*~K-Ar age reported for the geologic unit; ( ) is estimated from the stratigraphic sequence.
*2Total Fe as Fe203.
*3The errors are 2a of the mean.
320
in distilled water, and crushed in an agate mor-
tar. Major elements, Rb and Sr concentrations
were determined by Rigaku ® Symaltics 3530
and 3080 XRF (X-ray fluorescence) spectrom-
eters. Powdered samples of ~ 100 mg were dis-
solved and chemically separated for mass-
spectrometric anaysis. Nd was measured as the
metal species by a double Re filament with a
Finnigan ® MAT 261 E mass spectrometer at
Kyoto Sangyo University. The measured
~43Nd/144Nd ratios were normalized to
146Nd/144Nd of 0.7219 and trace contributions
from 144Sm were checked by monitoring 147Sm.
The average La Jolla standard value obtained
during the study was 143Nd/144Nd=0.511845
+0.000010 (n=7). The Nd isotopic ratio of
BCR-1 was0.512626+0.000031 (n=2). Srwas
measured on a double Re filament or a Ta single
filament. We obtained STSr/S~Sr ratios of
0.710220_+0.000014 (n=3) for NBS-987,
0.707992+0.000006 (n=4) for Eimer &
Amend ® SrCO3 and 0.704980+0.000018
(n=2) for BCR-I. The total blanks of the
chemical procedures were negligible.
Isotopic ratios are expressed in the notation
where end (0) is the measured (m) deviation in
parts in 104 of the 143Nd/~44Nd ratio from the
present-day chondritic value which is calcu-
lated to be 0.512622 by assuming that La Jolla
Nd standard ( 143Nd/144Nd = 0.511845) corre-
sponds to end:--15.15 (Wasserburg et al.,
1981 ). Similarly, esr (0) is the deviation in parts
in 104 of the measured STSr/S~Sr ratio from the
inferred reference value of 0.7045 correspond-
ing to a present-day modal undifferentiated
mantle reservoir (UR).
4. Results
Because of the young ages of the samples, no
corrections for in situ decay of Nd have been
made. Analytical results are listed in Table I.
Nd and Sr isotopic data are plotted on an or-
dinary eNd-esr correlation diagram in Fig. 2. Nd
and Sr isotopes show a substantial variation ranging from -0.8 to 8.3 and from -25.6 to
13.3 e-units, respectively. The most depleted
value is found in the Quaternary Kampuzan
lava which shows eNd of 8.3 and esr of -25.6
(data from Nohda and Waserburg, 1981). It
implies that the Kampuzan magma is derived
from a very depleted source similar to MORB
(mid-ocean ridge basalt). The most enriched
values are found in AKT2.3 ( eNd = -- 1, eSr= 13 )
and NS-428 (data from Nohda and Wasser-
burg, 1986). Except for samples with esr>0,
they plot approximately along the mantle ar-
ray. Fig. 3a and b shows the relationships be-
tween eNd and eSr and geologic age of the
samples. As pointed out by Nohda and Wasser-
burg (1986), Nd and Sr isotopic ratios are reg-
ular functions of geologic age; the volcanic
samples of the early Miocene show an enriched
isotopic signature, but a continuous shift to de-
pleted values is observed in the younger lavas.
It should be stressed that the present results are
not brought by geographic heterogeneity be-
cause the samples are obtained in a very limited
area of 30 × 60 km in the back-arc side of NE
Japan.
~Nd
8i
6
4
2
0 87Sr/8~r
0.7030 0,7040 0.7050
~v i i r i 1
~~ AKITA, NE JAPAN
I I I I l I I I
-20 -10 0 10
Fig. 2. Plots of end VS. es, of the volcanic rocks of the back-
arc side in NE Japan arc ( open circles = Tertiary volcanics;
solid circles = Quaternary volcanics; BE = bulk Earth). Data
from Nohda and Wasserburg (1981, 1986) are also plotted.
~Nd
8
87Sr
86Sr
0.705
0.704
0.703 [ i i [ I i i
-• ~ Nd Isotope Ratio
©
(a) 0 0
I I I I I I J
0 10 20 30
AGE (Ma)
©
b Sr Isotope Ratio
• )
I J i I I I
0 10 20
AGE (Ma) I
3O Es,
10
-10
-20
Fig. 3. Nd and Sr isotopic ratios vs. age for the volcanic
rocks from the back-arc side of the NE Japan arc. Symbols
are the same as those in Fig. 2. The isotopic compositions
show contrasting temporal variations from enriched to de-
pleted, in comparison with the constant isotopic composi-
tions of the volcanic rocks along the volcanic front of the
same arc (dotted areas). 321
ENd (a) PRE-OPENING
~'O.~
~Nd
I I •
100 km 50 V F. lO
~'Sr
0
-10
6
ENd
4 (b) POS T'OPENING
\ /Sr
l\ ,/
//O
////
/G ~Nd
1OOkm 40
Esr
-20
-30
Fig. 4. a. Nd and Sr isotopic profile across NE Japan at
~ 20 Ma during the pre-opening stage, estimated from Fig.
3a and b. Sr isotopic compositions increase towards the
back-arc side.
b. Nd and Sr isotopic profile across NE Japan along 40 ° N
of the Quaternary volcanics (Nohda and Wasserburg, 1981 ).
5. Discussion
The temporal variation of Nd and Sr isotopes
is synchronous with the opening of the Japan
Sea inferred from paleomagnetic evidence
(Otofuji et al., 1985a), magnetic anomalies in
the Japan Sea (Isezaki, 1986; Kono, 1986) and
a sharp change in marine fauna (Chinzei, 1986).
This synchronism indicates that there is a com-
mon cause for the opening of the Japan Sea and
the temporal variation of Nd and Sr isotopes of
the volcanics. During the opening of the Japan
Sea, the arc-trench system retreated from the
Asian continent toward the Pacific side, which
implies a possibility of compositional change of
the mantle wedge through tectonic transition.
The temporal variation of Nd and Sr isotopes
322
is considered in terms of the change of the me-
gatectonic structure related to the opening of
the Japan Sea basin.
5.1. Crustal contamination
The narrow linear array of the present data
on an Nd-eSr correlation diagram indicates a
mixing process between enriched and depleted
components in the magma source region
throughout the period. Nohda and Wasserburg
(1986) considered involvement of two end-
member components for magma generation: a
depleted component similar to a MORB-type
source (MORB source) and an enriched one
represented by the continental crust. The iso-
topic shift was then attributed to the change in
degree of crustal contamination and inferred
from the proportions of the contribution of
these two components due to a thinning pro-
cess of the continental crust accompanying the
back-arc opening of the Japan Sea.
Although crustal thinning seems to have been
a substantial process for the region around the
spreading center of the Japan Sea basin, the
present thickness of the crust is observed as
~ 30 km at around Akita (Yoshii, 1979) and is
equivalent to that of the Sikhote Alin and the
eastern part of the Asian continent (Soller et
al., 1982; Tan, 1987). These data suggest that
NE Japan had not been subjected to crustal
thinning. In addition, the ultramafic nodules
found at the Quaternary volcanics of Ichinom-
egata show a depleted Nd and Sr isotopic sig-
nature similar to the nearby Quaternary lavas
of the Kampuzan and Moriyoshiyama volcan-
oes (Nohda and Wasserburg, 1986). It may be
unrealistic to claim that only the Tertiary vol-
canics were subjected to crustal contamination.
Thus, we attribute the major governing factor
of the temporal variation of Nd and Sr isotopes
to a chemical transition of the upper mantle
during the opening of the Japan Sea. 5.2. Nature of the mantle wedge during the pre-
opening stage
We estimated the chemical structure of the
mantle wedge of the NE Japan region during
the pre-opening stage from the Nd and Sr iso-
topic data. We can make use of the unchanged
Nd and Sr isotopic data from the tholeiitic ba-
salts that are ranging from 22 Ma to the present
and found along the present volcanic front of
NE Japan (dotted area of Fig. 3a and b; Tat-
sumi et al., 1988 in this issue ). With these data
we can infer an isotopic profile of Nd and Sr
across the NE Japan region for the pre-opening
stage as shown in Fig. 4a which is similar to
those of the continental margin obtained at
California, U.S.A. (DePaolo, 1981a) and Pa-
tagonia, Argentina (Hawskesworth et al.,
1979), and also to the central section of Japan
(Nohda and Wasserburg, 1981). There, the
volcanics show a more enriched isotopic signa-
ture of Nd and Sr with increasing distance from
the trench. Such an isotopic profile seems to be
charateristic for the continental margin vol-
canic arc, and requires an explanation from the
viewpoint of the mantle structure.
The mantle wedge of the continental margin
is composed of sub-continental lithospere
( SCL ) which is chemically more enriched than
the MORB source (All~gre et al., 1982; Haw-
kesworth et al., 1983; Pearce, 1983). Before the
opening of the Japan Sea, the mantle wedge of
NE Japan was composed of SCL that was
thickened towards the back-arc side from the
trench side. During the subduction of the
oceanic plate, the depleted asthenosphere un-
derlay the SCL. Therefore, we propose the most
probable structure of the mantle wedge of NE
Japan for the pre-opening stage, as shown in
Fig. 5a. Such a two-layered structure of the up-
per mantle is characteristic of a continental arc.
A petrogenetic model is already discussed
elsewhere for the volcanics of NE Japan ( Tat-
sumi, 1986; Tatsumi et al., 1988 in this issue).
Here we emphasize that the Nd and Sr isotopic
signatures of the continental margin volcanic
arcs are mainly determined by the ratio of
thickness between SCL and the underlying
MORB source. The partially molten mantle
diapir would initiate within the depleted part of
the mantle wedge by addition of the enriched
fluid originally extracted from the downgoing
slab, then would upwell through the high-tem-
perature region ( > 1400°C) and release a pri-
are scattered with large variation but show gen-
erally higher Hf/La ratios (Ebihara et al.,
1984). The older volcanic rocks of the area show
higher La/Yb ratios, and lowering of the La/Yb
ratio with decreasing age is reported (Ebihara
et al., 1984). Such geochemical observations
give further confirmation of the enriched char-
acteristics of the mantle wedge of NE Japan
during the pre-opening stage.
5.3. Characteristics of the post-opening NE
Japan arc
The Nd and Sr isotopic profile for the post-
opening stage is represented by the Quaternary
volcanics across NE Japan (Nohda and Was-
serburg, 1981), and reproduced in Fig. 4b. It
shows a regular increase in eNd and decrease in
esr from the volcanic front to the back-arc side,
and has quite opposite sense with those of the
pre-opening and active continental margins. If
we suppose a homogeneous and depleted MORB
source for the mantle wedge of NE Japan, the
undepleted or less-depleted isotopic signature
of Nd and Sr from the volcanics of the trench
side strictly requires an involvement of the en-
riched crustal component. Sakuyama and Nes-
bitt (1986), however, deny such crustal
contamination to explain the chemical charac-
teristics of the basalts from the trench side, al-
though they were aware that their
interpretation did not match the isotopic char-
acteristics such as those in Fig. 4b. Masuda and
Aoki (1978) recognized that the Quarternary 323
tholeiites along the volcanic front and those of
the Izu-Marianas show differences in the REE
patterns. Their data suggest that the mantle
wedge of NE Japan is more enriched compared
to those of the intra-oceanic arc (A-type arc of
Nohda, 1984). Nevertheless, a depleted source
is required for the mantle of the back-arc side
to explain the depleted isotopic signature of the
volcanics from the back-arc side.
We propose a two-layered structure of the
mantle wedge of NE Japan to explain the ob-
tained Nd and Sr isotopic signature, as shown
mary basalt magma at a depth of ~ 30 km.
During upwelling, the mantle diapir was chem-
ically in open-system conditions, and had a
chance to react with the overlying SCL mantle
(DePaolo, 1981b). Consequently the Nd and Sr
isotopic compositions of the lavas are deter-
mined by the enriched fluid extracted from the
slab and the volume ratio between the SCL and
the MORB source within the mantle wedge.
The fluid extracted from the downgoing slab
is originally related to seawater and oceanic
sediments and is enriched in incompatible ele-
ments with larger ionic radii, but their isotopic
composition is not known (Tatsumi et al.,
1986). Seawater enriches Sr isotopes but is not
effective for Nd (DePaolo and Wasserburg,
1977). Except for the continental crust, oceanic
sediment is the most effective candidate to lower
the Nd isotopic composition in island-arc en-
vironments, but considerable amounts are re-
quired to alter the depleted Nd isotopic value to
an undepleted value (Nohda and Wasserburg,
1981). It is known that basalts of the intra-
oceanic arc ( A-type arc of Nohda, 1984) show
an average eNd of 8.1 which is lower than those
of N-type MORB ( end = 10). The mantle wedge
of an A-type arc is composed of the oceanic
mantle, the MORB source. Since the fluid from
the downgoing slab is the only possible com-
ponent to lower the end of lavas, we can thus
estimate the degree of the fluid effect to the
mantle wedge as being 1.9 e-unit lowering of Nd
324
~ /(C_~ (a) PRE-OPENING
j
SCL ~ d~.~j ~
: .......... j_:~ BTKDO~fNOFAMPHtSOLE&CHLORIT£
;>
Ja0an S ...... ~ ~.-(C)POST-OPENING
--i i
Fig. 5. A sketch for the magma genesis in the Neogene NE
Japan arc. The subducted lithosphere can supply HzO to
the mantle wedge only beneath the fore-arc region. The slab-
derived H20 reacts with mantle wedge materials to form
hydrated peridotite, which is dragged downwards on the
slab. H20 released through the breakdown of hydrous phases
in the dragged hydrated peridotite causes partial melting
in the mantle wedge beneath the volcanic arc. The sub-
continental lithosphere ( SCL; dotted area) becomes thin-
ner beneath the back-arc side of the NE Japan arc through
the injection of asthenosphere with composition similar to
the depleted MORB source during the back-arc spreading.
G T. represents an isotherm of ~ 1000 ° (Tatsumi, 1986).
isotopes. The average end of the trench side ba-
salts of NE Japan is 2.8, which means that the
original mantle wedge ( SCL mantle + MORB
source) had a end of 4.7 [2.8--(--1.9)=4.7]
before the fluid addition. The Nd and Sr iso-
topic compositions of the SCL are not arbitrar-
ily determined from this value because the Nd and Sr contents in the SCL are not known. But
we can conclude that the SCL of NE Japan is
characterized by the enriched isotopic signature.
The enriched signature of the SCL of NE Ja-
pan is also supported by trace-element data.
Tertiary volcanics of NE Japan define a con-
stant Hf/La ratio, while Quaternary volcanics
in Fig. 5c. The SCL is thicker beneath the vol-
canic front and gradually thins towards the
back-arc side. The mechanism of magma gen-
eration for the post-opening stage is similar to
that of the pre-opening stage. Namely, during
upwelling of the partial molten mantle diapir
within the higher-temperature zone
( > 1400 °C), chemical reaction with the over-
lying mantle determines the Nd and Sr isotopes
of the magma. The magmas of the back-arc side
require a higher involvement of the MORB
source to match their depleted isotopic signa-
tures of Nd and Sr. The trench side basalts show
an unchanged and undepleted isotopic signa-
ture of Nd and Sr through 22 Ma, which implies
that the ratio of the SCL to the MORB source
concerning the magma generation was kept
constant and higher than those for the back-arc
side of the post opening. In addition, a linear
temporal variation of Nd and Sr isotopes is a
result of an identical Nd/Sr ratio between two
end components in the mixing process (De-
Paolo and Wasserburg, 1979).
It is well recognized that incompatible ele-
ments of lavas from NE Japan behave in a sim-
ilar manner to K20 which increases with
increasing distance from the trench (Fujitani
and Masuda, 1981; Gill, 1981; Sakuyama and
Nesbitt, 1986). This geochemical feature is at-
tributed to a decrease in degree of partial melt-
ing of a homogeneous source in the same
direction ( Sakuyama and Nesbitt, 1986). This
is not in conflict with our present model of the
two-layered mantle, because the high abun-
dance of incompatible elements of the lavas
from the back-arc side is explained by the small
degree of partial melting of the source, predom-
inantly MORB source, to match with the de-
pleted isotopic signature. The flat pattern of
REE abundances found for the basalts from the
trench side is possibly caused by a higher degree
of partial melting of the mantle which involves
a higher volume of SCL to have the undepleted
isotopic composition of Nd and Sr.
5.4. Opening of the Japan Sea; asthenospheric
injection
During the pre-opening stage, the volcanics
had a more enriched isotopic signature of Nd
and Sr with increasing distance from the trench,
which was caused by a higher involvement of
the SCL for magma generation in the region of
the back-arc side than those in the trench side.
With decreasing age, Nd and Sr isotopes of the
lavas from the back-arc side shift to the de-
pleted side, which implies a change in the vol-
ume ratio between SCL and MORB source. This
structural transition - the thinning of SCL and
complementary growth of the depleted MORB
source in the mantle wedge of the back-arc side
- took place along with the opening of the Ja-
pan Sea.
We believe that the asthenosphere was in-
jected into the mantle wedge of the back-arc side
and caused growth of the depleted mantle
(MORB source). Moreover, the astheno-
spheric injection caused the opening of the Ja-
pan Sea. Due to injection of the asthenosphere,
SCL was subjected to the thinning process. The
thinning of SCL gave tensional stress to the
overlying continental lithosphere, that finally
rifted. Subsequent migration of continental
fragments as the Japanese islands is traced by
paleomagnetic data. The idea of astheno-
spheric injection is supported by eastward as-
thenospheric flow which explains the
distribution of the back-arc basins in the west-
ern Pacific (BostrSm, 1981), and may be re-
lated with migration of the hot region of
Miyashiro (1986). 325
Models explaining processes during back-arc
opening are discussed by Taylor and Karner
(1983). Diapirism and the return flow within
the enriched SCL could not supply the depleted
magmas and fail to explain the linear trend of
Nd and Sr isotopes observed in the volcanics
from the back-arc side of NE Japan. It is prob-
able that induced flow is insufficient to initiate
back-arc spreading where the overlying litho-
sphere is continental (ToksSz and Bird, 1977).
Trench retreat seems to be present in every
subduction zone, but cannot explain the spatial
distribution of back-arc basins in the western
Pacific. Moreover, trench retreating necessar-
ily involves a volume deficit beneath the region
of the back-arc side and requires injection of
the asthenosphere.
The mantle wedge beneath the present vol-
canic front was not under the influence of the
asthenospheric injection. It shows an un-
changed and undepleted Nd and Sr isotopic sig-
nature through the period from 22 Ma, which is
correlated with the observation that the inter-
mediate Q and V layers are observed in the
mantle wedge between the volcanic front and
the aseismic front (Matsuzawa et al., 1986).
Thus, the asthenospheric injection seems to
match the isotopic trend of Nd and Sr found in
the volcanic rocks and the opening of the Japan
Sea from paleomagnetic evidence.
Acknowledgements
The first author (S.N.) acknowledges Pro-
fessor G.J. Wasserburg for the opportunity to
start research on Nd and Sr isotopes of the vol-
canic rocks of NE Japan. He also thanks the
Ministry of Education, Science and Culture of
Japan, and Kyoto Sangyo University for their
support in setting up the laboratory for isotope
geology. Dr. Murata of Rigaku Corporation
provided us with the opportunity to use their
XRF facilities. Professor T.N. Robb, Dr. T.
326
Hasenaka and Miss Y. Maeda read the manu-
script and offered many helpful improvements.
This work was supported by the Asahi Scholas-
tic Fund, 1986, the Inamori Foundation, and a
Grant-in-aid for Special Project Research (No.
62103007) from the Ministry of Education,
Science and Culture of Japan.
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Available online at www.sciencedirect.com SCIENCE DIRECT? ELSEVIER Gondwana Research 9 (2006) 142 - 151 Provenance of Paleozoic-Mesozoic sedimentary rocks in the Inner Zone of Southwest Japan: An evaluation based on Nd model ages 1 1e 1 11 1 q o *e T. Hamamoto f, Y. Hayasaka , Y. Ikeda h, M. Yuhara ′, Y. Tainosho j a Graduate School of Science and Technology, Niigata University, Nigata 950-2181, Japan b Faculty of Culture and Education, Saga University, Saga 840-8502, Japan Nittoc Construction Co. Ltd.,Yamaguchi Branchi, Hirano 1940-16, Miyanoshita, Yamaguchi 753-0011, Japan dYachiyo Engineering Co.Ltd.,Naka-Meguro 1-10-23, Tokyo 153-8639, Japan 200km f Dia Consultants Co. Ltd., Kanayama-cho 1-6-12, Nagoya 456-0002, Japan Graduate School of Science, Hiroshima University, Hiroshima 739-8526, Japan 13 iFaculty of Science,Fukuoka University,Fukuoka 814-1180,Japan Department of Natural Environment, Kobe University, Kobe 657-0011, Japan Available online 910January 2006 Abstract Nd model ages using depleted mantle (TDM) values for the sedimentary rocks in the Inner Zone of the SW Japan and western area of Tanakura Tectonic Line in the NE Japan allow classification into five categories: 2.6-2.45, 2.3-2.05, 1.9-1.55, 1.45-1.25, and 1.2-0.85 Ga. The the source rocks in East China and U-Pb zircon ages. The provenance of 2.6-1.8 Ga rocks, which are reported from Hida-Oki and Renge belts and Kamiaso conglomerates, is inferred to be the Sino-Korean Craton (SKC). The 2.3-1.55 Ga rocks, mostly from Ryoke, Mino and Ashio belts, are originally related with the SKC and/or Yangtze Craton (YC). The provenances of the sedimentary rocks with 1.45 -0.85 Ga, from the Suo belt, Higo and some districts in the Mino and Ashio belts, are different from the SKC and YC. Especially, the Higo with 1.2-0.85 Ga is considered as a fragment of collision zone in East China. Akiyoshi belt probably belongs to the youngest age category of 1.2-0.85 Ga. probably derived from mafic igneous rocks within the Ryoke belt itself and from the adjacent Tamba belt. 2005 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. Keywords: Southwest Japan Arc; Paleozoic -Mesozoic sedimentary rocks; Nd model age; Sino-Korean Craton; Yangtze Craton 1. Introduction 1998; Fig. 1). The high P/T metamorphic terrane is subdivided into two belts namely Renge and Suo that are again classified as The Japan Arc was formed by a stepwise accretion from the western and eastern Suo. The low P/T metamorphic terrane continent to the Pacific Ocean side of the eastern margin of the includes Hida-Oki and Ryoke belts. The accretionary terrane is Eurasian Continent (e.g., Ichikawa, 1990; Isozaki and Mar- 15 uyama, 1991). This arc was separated from the continent, which Ashio belt in NE Japan is generally accepted as a northeastern was caused by opening of the Japan Sea basin in Early to Middle N Miocene (e.g., Otofuji et al., 1985). The Inner Zone of SW Japan lithostratigraphy indicates that the Inner Zone except Hida- Oki is divided into three terranes as high P/T metamorphic, low P/T metamorphic and accretionary complex (Nishimura, 1990, 267 (Fig. 2). The Hida-Oki belt has been considered to be formed in * Corresponding author. and around Sino-Korean Craton (e.g., Isozaki, 1997; Ishiwatari E-mail address: Kagami@gs.nigata-u.ac.jp (H. Kagami) and Tsujimori, 2003) (hereafter abbreviated as SKC). The Renge 1342-937X/$ - see front matter ? 2005 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2005.11.001 H.Kagamietal./GondwanaResearch9(2006)142-151 143 HighP/Tmet orphicterrar Akiyoshi belt laizurubelt LowP/ phicterra Hida-Okibelt oelt TT Hig ISTL Tsul Honshu Ryok RKyok TTL:TanakuraTectonicLine ISTL:Itoigawa-Shizuoka Tectonic Line 100km MTL:MedianTectonicLine Fig. 1. Generalized geological map of the Inner Zone of Southwest Japan Arc (after Ichikawa, 1990; Nishimura, 1998) and frequencies of TDMs. The extent of the Higo is after Hayasaka and Hirose (1997). Data sources; Table 1 (this study), Yamana (1988), Morris and Kagami (1989), Asano et al. (1990), Kagami et al. (1992), Tanaka (1992), Shimizu et al. (1996), Owada et al. (1999), Rezanov et al. (1999), Yuhara et al. (2000). The name of places enclosed by rectangles shows the sampling cites of analyzed samples. belt has been accreted to the Hida-Oki belt though its derived from these, collected from the main belts shown in Fig. accretional age is unknown. The Akiyoshi, Suo and Maizuru 1. The extraction procedures of Sm and Nd from rock powders belts have been accreted through the Middle to Late Permian were carried out following procedures given in Kagami et al. against the Hida-Oki and Renge belts. The Mino-Tamba and (1989). The isotopic analyses were performed on MAT261 Ashio belts and Ryoke belt were accreted in Jurassic against the and MAT262 mass spectrometers at Niigata University. composite block which was formed by pre-Jurassic accretion 6I7L'0=PN++1/PN9t1 0 pazeou aM s0ne1 PNt+1/PNet1 (Nishimura, 1998). and are reported relative to 143Nd/144Nd=0.512115 for Ndi-1 Initial formation ages of the precursors of sedimentary (Standard of Geological Survey of Japan) corresponding to 143Nd/144Nd=0.511858 of LaJolla (Tanaka et al., 2000). Mean rocks and metamorphic rocks can be inferred by Nd model ages with respect to the depleted mantle (DM) (e.g., uncertainty of 143Nd /144Nd ratio for each sample is ca. 0.003% DePaolo, 1988; Milisenda et al., 1988; Chen and Yang, as 2o-value. Sm and Nd concentrations were obtained using 2000). We estimate the initial formation ages of precursors for Paleozoic to Mesozoic coarse-grained sedimentary rocks ratios are 0.1% as standard deviation. TDMs were calculated and metamorphic rocks that prevail in the Inner Zone of the =+ 9010=PN/S0141S0=PN1/PN1 SW Japan Arc. The Nd model age data provide significant insights into the history of a mature island arc like the SW 6.54 × 10- 12y-1. TDMs for the samples were calculated with Japan Arc. ratio less than 0.13. The Sm and Nd isotopic ratios and 2. Samples and analytical procedure TDMs are shown in Table 1. Additional 62 data from published literatures were also used for TDM, the source of The 72 samples analyzed in this study comprise argillaceous PNp+/PNti u 'I biH Jo uondeo s u pasl s! uoum to psammitic sedimentary rocks and the metamorphic rocks data were corrected using values of standard samples LaJolla 144 H. Kagami et al. / Gondwana Research 9 (2006) 142-151 HighP/T LowP/T Accretionaryterrane 3.2.LowP/T metamorphic terrane neta.terra neta.terra. (1) Hida-Oki belt; 2.6-1.8 Ga. Metasedimentary rocks are O 1 distributed in Oki-Dogo island with TDMs of 2.55- 2.45 Ga. These ages are older than Nd model ages compared with those of CHUR (2.3-2.2 Ga; Tanaka (Ma and Hoshino, 1987), and probably coincide with TDMs Cretaceous if the values are recalculated using DM. However they 135 L are older than Sm-Nd whole rocks isochron ages of Jurassic M 1.98 and 1.96 Ga (Tanaka and Hoshino, 1987) and U- Pb zircon ages of 1.96 Ga (Yamashita and Yanagi, E 205 1994). Eight Sm-Nd isotopic compositions of metase- L Triassic dimentary rocks from Hida in Chubu district were 250 reported by Asano et al. (1990) and Tanaka (1992). Permian Four samples have TDM of 2.1-1.8 Ga. Other three samples have extremely high 143Nd/144Nd (>0.5126) Carboni- ratios. These samples plot between 0.5 and 1.0 Ga lines ferous shown in Fig. 3 and are characterized by low Zr/Ti and 355 SiO2/AlzO3 ratios as pointed out by Kawano et al. L Devonian W (2006), which strongly suggest that they are mainly E composed of immature clastic materials derived from 410 Silurian Sandstone&Mudstone basaltic to andesitic volcanic rocks (Musashino, 1992). 43 Limestone Another sample has high 147Sm/144Nd ratio (>0.14) L Chert & Sillieous rocks though its 143Nd/144Nd ratio is lower than 0.5125. Ordovician M Mafic volcanics (2) Ryoke belt: Ryoke belt preserves a wide age range E 510 between 1.9 and 0.85 Ga. However, each district within Cambrian Plutonism this belt has distinct TDMs. Higo district is located in 570 600 Metamorphis sm（h,m,l:high Preterozoic 1100 + Kyushu (Fig. 1) and it is uncertain whether the Higo is 1600~1800 western extension of the Ryoke belt or not. TDMs range Fig. 2. Generalized and simplified lithostratigraphy of the Inner Zone of SW between 1.25 and 0.85 Ga of the Higo highly deviate Japan Arc. Modified from Fig. 2 of Ichikawa (1990). Age of metamorphism of from 1.9 to 1.55 Ga of the typical Ryoke metasedimen- the Suo belt is after Nishimura (1998). tary rocks as described below. The Higo metasedimen- tary rocks contain inherited Proterozoic zircons with ages (0.511858), BCR-1 (USGS standard, 0.512638, Kagami et of 1.78 Ga (Osanai, unpublished data). U-Pb chemical al., 1989), JB-1a (GSJ standard, 0.512784, Kagami et al., dating of monazites from this belt using the CHIME 1989), and JNdi-1(GSJ standard, 0.512115, Tanaka et al., method (Suzuki et al.,1998) shows that the Proterozoic 2000). ages are scattered between 1.9 and 0.8 Ga, but most of them concentrate between 1.4 and 0.8 Ga. 3. Nd model ages of sedimentary rocks in the Inner Zone, SW Japan Arc Yanai-Uojima district in the Ryoke belt yields ages of 1.8-1.5 and 1.3 Ga. One sample (sample no. 27) with It is noteworthy that TDMs obtained from certain belt (or 1.3 Ga was collected from the western most segment of district in the belt) of the terranes are not scattered at random Yanai-Uojima district. Some Ryoke granites occurring but cluster around restricted time range as shown in TDMs in Yanai district contain granulite xenoliths which are frequency (Table 1, Fig. 1). We thus recognize distinct age o nn 143Nd/144Nd=0.5123~0.5125, Table 1) populations as discussed in the sections below. (present (Hamamoto et al., 1999). Accordingly, the sample with 3.1.HighP/T metamorphic terrane 1.3 Ga age is probably related to the metasedimentary rocks from Higo district. TDM of Shishijima rocks show (1) Renge belt; 1.83 Ga, 1.4 Ga, (2) western Suo belt; 1.2 Ga (sample no. 30), which is quite different from 1.45-1.2 Ga, (3) eastern Suo belt; 1.25-1.1 Ga. One those of typical Ryoke belt. Age data could be not sample (no. 1) of the Renge belt gives an old age of computed from another sample (no. 31) due to high 143Nd/144Nd ratio of 0.512729. Rocks from the Kinki 1.83 Ga, and another 1.4 Ga. Suo psammitic schist from Kyushu contains detrital zircon with an age of 1.91 district has normal TDM as the Ryoke belt ranging (Miyamoto and Yanagi, 1996), 1.75 and 1.4-1.3 Ga between 1.8 and 1.55 Ga without any exception. (Tsutsumi et al., 2000), and reset of the ages are Although we analyzed Sm and Nd isotopic compositions coincident with TDMs. for four metasedimentary rocks from Mie district, TDM H.Kagamiet al./Gondwana Research9 (2006)142-151 145 Table 1 147Sm /14Nd ratios and 143Nd /144Nd ratios and TDMs Name 147Sm/ Number of samples 143Nd/ Name of belts TDM Name Locations of samples 144Nd 144Nd and districts of rocks (Ga) of samples 1. High P/T metamorphic terrane Renge belt Schist* 0.511980 1 0.1166 1.83 306-6 Higashinagano, Toyota-cho, Yamaguchi P. 2 Schist* 0.1305 0.512173 306-8 Toyoga-dake, Toyota-cho, Yamaguchi P. 3 Suo belt (western) Gneiss* 0.1186 0.512399 1.20 HA24R Shimoirahara, Saigawa-machi, Fukuoka P. 4 Gneiss* 0.1270 0.512500 HA45 Kamirahara, Saigawa-machi, Fukuoka P. 5 Gneiss* 0.1512 0.512435 IM57 Tsuno, Soeda-machi, Fukuoka, P. 6 Schist* 0.1248 0.512309 1.45 AK8 Oomukai, Tokuyama-shi, Yamaguchi P. 7 Schist* 0.1212 0.512311 1.38 AK9 Mitake, Tokuyama-shi, Yamaguchi P. Schist* 0.512411 8 0.1282 1.32 AK14 Deai, Nishiki-cho, Yamaguchi P. Schist* 9 Suo belt (eastern) 0.1230 0.512420 1.23 1005 Tari, Nichinan-cho, Tottori P. 10 Schist* 0.1209 0.512401 1.23 1007 Tari, Nichinan-cho, Tottori P. 11 Schist* 0.1235 0.512494 1.11 1008 Tari, Nichinan-cho, Tottori P. 2. Low P/T metamorphic terrane 12 Higo district Gneiss* 0.1116 0.512309 1.26 Grt-Crd-Bt Gn-1 Tanohira, Toyono-cho, Kumamoato P. 13 Gneiss* 0.1113 0.512514 Grt-Crd-Bt Gn-2 Tanohira, Toyono-cho, Kumamoato P. 14 Gneiss* 0.1052 0.512504 Grt-Crd-Bt Gn-3 Tanohira, Toyono-cho, Kumamoato P. 15 Gneiss* 0.1169 0.512466 1.08 Grt-Bt Gn-1 Tanohira, Toyono-cho, Kumamoato P. 16 Gneiss* 0.1169 0.512426 1.14 HG-01 Uchida, Matsubase-cho, Kumamoto P. 17 Gneiss* 0.1309 0.512463 AM-01 Takado, Ryugatake-cho, Kumamoto P. 18 Granulite* 0.1120 0.512476 1.01 71602A-a Uchida, Matsubase-cho, Kumamoto P 19 Granulite* 0.0955 0.512495 0.85 Spr-Co Granulite-1 Uchida, Matsubase-cho, Kumamoto P 20 Granulite* 0.1308 0.512459 Spr-Co Granulite-2 Uchida, Matsubase-cho, Kumamoto P 21 Granulite* 0.1233 0.512476 1.14 Spr-Co Granulite-3 Uchida, Matsubase-cho, Kumamoto P 22 Granulite* 0.1275 0.512470 1.20 71602A-e Uchida, Matsubase-cho, Kumamoto P 23 Granulite* 71602A-g 0.1066 0.512462 0.98 Uchida, Matsubase-cho, Kumamoto P 24 Granulite* 0.1417 0.512521 71602A-h Uchida, Matsubase-cho, Kumamoto P 25 Granulite* 0.1326 0.512454 Spr-Co Granulite-4 Uchida, Matsubase-cho, Kumamoto P 26 Ryoke belt Gneiss* 0.1853 0.512621 N08331911 Himuro-dake, Shuto-cho, Yamaguchi P. (Yanai district) 27 Gneiss* 0.1144 0.512309 1.29 TN0903224 Kajitori-misaki, Hikari-shi, Yamaguchi P. 28 Gneiss* 0.1177 0.511923 1.94 TN646151 Himi-misaki, Ooshima-cho, Yamaguchi P. 29 Ryoke belt (Uojima) Gneiss* 0.0977 0.511792 1.78 TN809422 Enoshima, Uoshima-son, Ehime P. 30 Ryoke belt (Shishijima) Gneiss* 0.1199 0.512729 KO79072801 Shishijima, Takuma-cho, Kagawa P. 31 Gneiss* 0.1215 0.512432 1.19 MA78073006 Shishijima, Takuma-cho, Kagawa P. 32 Ryoke belt Mudstone 0.1091 0.511977 1.71 1083-21 Kannabi-yama, Kyotanabe-shi, Kyoto-fu (Kinki district) 33 Mudstone 0.1250 0.512230 1.58 98090707 Kannabi-yama, Kyotanabe-shi, Kyoto-fu 34 Shale 0.1090 0.512074 1.56 98090706 Kannabi-yama, Kyotanabe-shi, Kyoto-fu 35 Shale 0.1109 0.511940 1.79 98090703 Kannabi-yama, Kyotanabe-shi, Kyoto-fu 36 Ryoke belt Gneiss* 0.1605 0.512866 Bano-gawa 一 upper reaches of Bano-gawa, Iga-shi, Mie P. (Mie district) 37 Gneiss* 0.1681 0.512272 14-22T Minami-nagano-gawa, Misato-mura, Mie P. 一 38 Gneiss* 0.1304 0.512405 07-10 north of Kasatori-yama, Iga-shi, Mie P. 39 Gneiss** 0.1572 0.512416 一 7-101 north of Kasatori-yama, Iga-shi, Mie P. 3.Accretionalterrane 40 Akiyoshi belt Mudstone 0.1391 0.512729 MGT Mugi-tani, Fukue-mura, Yamaguchi P. 41 Shale 0.1205 0.512547 921 Nishidera, Mine-shi, Yamaguchi P. 42 Shale 0.1517 0.512780 994 Tsujinami, Misumi-cho, Yamaguchi P. 43 Shale 0.1514 0.512847 996 Tsujinami, Misumi-cho, Yamaguchi P. 44 Tamba belt Mudstone 0.1133 0.512256 1.36 AK1 Miune-gawa, Nichihara-cho, Shimane P. (Chugoku district) 45 Mudstone 0.1191 0.512149 1.61 AK5 Miune-gawa, Nichihara-cho, Shimane P. 46 Mino belt (Kiso) Sandstone 0.1061 0.511591 2.20 98102201 Ogiso, Kiso-mura, Nagano, P. 47 Sandstone 0.1064 0.511626 2.16 98102202 Oku-Kiso-ko, Kiso-mura, Nagano P. 48 Sandstone 0.1073 0.511563 2.30 02110203 Kusari-gawa, Asahi-mura, Nagano P. 49 Sandstone 0.1119 0.511633 2.27 02110204 Kusari-gawa, Asahi-mura, Nagano P. 50 Slate 0.1100 0.512056 1.61 98102203 Narai-dam, Narakawa-mura, Nagano P. 51 Sandstone 0.1171 0.512148 1.58 98102204 Gonbei-toge, Minamiminowa-mura, Nagano P. 52 Sandstone 0.1072 0.511985 1.67 98102205 Kasagadaira, Narakawa-mura, Nagano P. 53 Slate 0.1027 0.511972 1.63 02110201 Yokokawa-keikoku, Tatsuno-machi, Nagano P. 54 Slate 0.1063 0.511838 1.86 02110202 Yokokawa-keikoku, Tatsuno-machi, Nagano P. (continued on next page) 146 H. Kagami et al. / Gondwana Research 9 (2006) 142-151 Table 1 (continued) Name 147Sm/ 143Nd/ Number of samples Name of belts TDM Name Locations of samples 144Nd 144Nd and districts of rocks (Ga) of samples 3.Accretional terrane 55 Slate 0.1079 0.511792 1.98 02110301 Shima-shima-dani, Azumi-mura, Nagano P. 56 Sandstone 0.1051 0.511736 2.00 02110302 Shima-shima-dani, Azumi-mura, Nagano P. 57 Slate 0.1119 0.512076 1.62 02110305 Kurosawa-dam, Misato-mura, Nagano P. 58 Slate 0.1327 0.512113 02110306 Susado, Hotaka-machi, Nagano P. 59 Slate 0.1013 0.512287 1.18 02110303 Sawando, Azumi-mura, Nagano P. 60 Slate 0.1092 0.512233 1.35 02110304 Sawando, Azumi-mura, Nagano P. 61 Ashio belt Sandstone 0.1212 0.512026 1.85 00081004 Matou, Ashio-machi, Tochigi P. (southern Ashio) 62 Sandstone 0.1094 0.512038 1.62 01092604 Kusagi, Azuma-mura, Gunma P. 63 Ashio belt (Tsukuba) Gneiss* 0.1244 0.512085 1.81 98081602 Busshouji, Yasato-machi, Ibaraki P. 64 Gneiss* 0.1193 0.512069 1.75 98081605 Hirasawa, Tsukuba-shi, Ibaraki P. 65 Mudstone 0.1126 0.512047 1.66 98081702 Miyagasaki, Yasato-machi, Ibaraki P. 66 Sandstone 0.1542 0.512163 98081703 Kamigou, Iwama-machi, Ibaraki P. 67 Ashio belt Sandstone 0.1091 0.511602 2.25 98081708 Tomiya, Iwase-machi, Ibaraki P. (southern Yamizo) 68 Mudstone 0.1047 0.511556 2.22 01092404 Oouchi, Batou-machi, Tochigi P. 69 Sandstone 0.1077 0.511605 2.21 01092406 Tomiya, Iwase-machi, Ibaraki P. 70 Mudstone 0.1064 0.511708 2.04 01092501 Kadoke, Iwase-machi, Ibaraki P. 71 Ashio belt Gneiss* 0.0931 0.512076 1.36 00080801B Kamehisa, Kurobane-machi, Tochigi P. (northern Yamizo) 72 Mudstone 0.0581 0.511993 1.13 00080801E Kamehisa, Kurobane-machi, Tochigi P. *; argillaceous, **; psammitic, P.; Prefecture. nos. 46-49; Miso-gawa complex, nos. 50-54; Kyogatake complex, nos. 55-58; Shima-shima complex, nos. 59 and 60; Sawando complex. The names of complexes are after Takeuchi et al. (1998). could not be obtained because their 143Nd / 144Nd 3.3. Accretionary terrane ratios exceed 0.5125 and/or greater than 0.13 (Table 1). Especially, three values (1) Akiyoshi belt: Even though four samples were analyzed, with 147Sm/144Nd>0.157 are rather abnormal for the we could not calculate TDMs because of extremely high sedimentary rocks occurring in the SW Japan Arc (Fig. 3). We discuss the source for the sedimentary (Table 1). These samples plot between 0.5 and 1.0 Ga rocks from Mie as well as two samples with lines shown in Fig. 3, and are characterized by low Zr/Ti 143Nd/144Nd ratio of 0.512621 (no. 26） from Yanai and SiO2/Al2O3 ratios, especially samples nos. 40, 42 and 0.512729 (no. 31) from Shishijima. TDMs of and 43 (refer to Table 1 in Kawano et al., 2006). Their thirty-one samples reported from Chubu district range Sm-Nd isotopic and chemical characteristics may be from 1.9 to 1.55 Ga. related to basaltic volcanic rocks such as HIMU (Tatsumi 0.5130 High P/T metamorphic terrane Renge 0.5Ga ■ Western Suo 1.5 Ga Easterm Suo 0.5125 O orphicterrane 1.0Ga XHida 口 !O+ 03H O 0.5120 Ryoke 由 田 田 3.5G Akiyoshi 2.0 Ga Tamba 0.5115 田Kamiaso Mino Kiso 蜀田 Ashio 2.5 Ga √ Tsukuba 0.5110 80.0 0.10 0.12 0.14 0.18 Yamizo 0.16 147sm/144Nd Fig. 3. 143Nd/144Nd vs. 147Sm/144Nd ratios relationship for Paleozoic to Mesozoic sedimentary rocks from SW Japan Arc. Sources for the Sm and Nd isotopic data are given in caption of Fig. 1. Stippled field of Yangtze Craton and 3.6-3.2 Ga igneous rocks in the world were cited from Chen and Yang (2000). Field of Sino- Korean Craton is defined using the data of Huang et al. (1986), Lan et al. (1995), Cheong and Chang (1997), Inomata (1999) and Lee et al. (2001). Straight lines line indicating 2.0 Ga. H. Kagami et al. / Gondwana Research 9 (2006) 142-151 147 et al.,2000) or rocks of tholeiitic to alkalic nature coincident with those from the complex of Chubu (Tazaki et al., 1989; Sano et al., 2000) of the Akiyoshi district in the Mino belt. greenstones. Psammitic rocks from Akiyoshi belt indicate U-Pb zircon ages of 0.3 Ga with inherited 4. Provenances of sedimentary rocks with old TDMs zircons of 2.7-2.4 Ga (Tsutsumi et al., 2000) (2) Maizuru belt: We did not analyze Sm and Nd isotopic As observed in this study, the TDMs obtained from various compositions of rocks from the Maizuru belt because, belts within the different terrains are not scattered at random this belt largely comprises volcanics and related sedi- but cluster around restricted time range, classified under mentary rocks which were formed in oceanic and island following age categories: (1) 2.6-2.45 Ga, (2) 2.3-2.05 Ga, arc settings during the Permian Period (e.g., Ishiwatari, (3) 1.9-1.55 Ga, (4) 1.45-1.25 Ga, and (5) 1.20-0.85 Ga. 1985a, 1985b; Koide et al., 1987; Sano, 1992; Herzig et (1) 2.6-2.45 Ga: Metasedimentary rocks from the Oki al., 1997). island (Hida-Oki belt) and granitic conglomerates from the (3) Mino-Tamba belt: Tamba belt constitutes the western Kamiaso district (Mino belt) belong to this age category part of the Mino- Tamba belt and is situated at Chugoku (Fig. 1). district of the SW Japan Arc. The rocks collected from Initial Nd isotopic ratios of some igneous rocks with activity this belt have TDMs of 1.6 and 1.35 Ga, the latter age ages from Middle Archean (ca.3.6 Ga) to Middle Proterozoic coinciding with the data from western part (Kyushu, (ca.1.5 Ga) occurring in the SKC are plotted near the Nd Chugoku district) of the SW Japan Arc, but distinct from evolution line of DM (Fig. 4, Shimizu et al., 1996; Lee et al., the normal TDM, 1.9-1.55 Ga, of the Ryoke belt. The 2001). However, most of the Late Archean (ca. 2.6 Ga) to Mino belt represents the eastern part of the Mino- Neoproterozoic (ca. 0.8 Ga) granitoid rocks and metasedimen- Tamba belt. TDMs of Kiso in the Chubu district are tary rocks from the SKC are plotted on the Nd evolution line divided into three groups of 2.3-2.15 Ga (nos. 46-49; Miso-gawa complex), 2.0-1.6 Ga (nos. 50-58; Kyoga- (0 Ga)= - 22) with 0.509572 (2.6 Ga, corresponding to eNd take and Shima-shima complexes) and ca. 1.3 Ga (nos. (2.6 Ga)=+6) (Shimizu et al.， 1996; Lee et al., 2001). 59, 60, Sawando complex). Thus, the complex having PN3) 8lIS'0 Jo one PNti/PNeti usad ou 1aMH the ages of 2.3-2.15 Ga is different from those of 2.0- (0 Ga)=-27.5) are more comparable with those from the 1.6 Ga and ca. 1.3 Ga (Takeuchi et al., 1998). These northern area of Korean Peninsula (Fig. 4) (Inomata, 1999). complexes are in contact separated by large faults. The s-values were calculated using the following CHUR's TDMs of 2.0-1.6 Ga ages overlap those of Ryoke 143Nd/144Nd (0 Ga)=0.512638, 147Sm/144Nd parameters: metasedimentary rocks from Chubu district. Psammitic (0 Ga)=0.1966. This evolution line with an inclination of rocks from Kamiaso in the Chubu district contain 97 N Jo y s1 (PN/S=) 0010 10 tI10 147Sm/144Nd value is calculated using conglomerates (mainly granitoid rocks) with Rb-Sr Ga. The former 143Nd/144Nd (0 Ga) of 0.511510 and the latter 0.511228, whole rock isochron ages of 2.06-1.89 Ga (Shibata and Adachi, 1974) and Sm-Nd whole rock isochron age of 2.07 Ga (Shimizu et al., 1996). TDMs of the (TDM= 2.6 Ga) from Kamiaso conglomerate and metamorphic conglomerates concentrate at 2.6 Ga. rocks (TDM=2.55-2.45 Ga) from Oki-Dogo island are also (4) Ashio belt: Although the Ashio belt is situated at the plotted on this line (Fig. 4). These data suggest that the age of NE Japan Arc, it forms the northeastern extension of 2.6 Ga marks one of most important periods of granitic crust the Mino-Tamba belt of SW Japan Arc. We analyzed formation in the SKC. This inference is based on the observation that the 147Sm/144Nd ratio of 0.114 or 0.100, the samples collected from Ashio, Tsukuba and Yamizo districts (Fig.1). Rezanov et al．(1999) analyzed Sm and Nd isotopic compositions of sedi- +15 mentary rocks of the northern district of Ashio belt PN3 Gneiss whose TDMs are between 1.45 and 1.6 Ga. Southern Granite +10 Ashio district has 1.85 and 1.6 Ga, which are slightly DM OAmphibolite Granulite older than the northern district. We obtained three +5 TDMs in four analyzed samples from Tsukuba, which Oki口, CHUR are 1.8-1.65 Ga. TDMs of four samples from southern Yamizo are older and range between 2.25 and 2.0 Ga. -5 Samples collected from northern Yamizo show ages of 冶 1.36 and 1.13 Ga, which are quite different from those -10 of southern Yamizo, and furthermore, they have Age (Ma) -27.5 -15 unusual TDM as the Ashio belt. Thus, the Ashio belt 500 10001500200025003000350040004500 has variable TDMs between 2.25 and 1.1 Ga. Fig. 4. eNd vs. Age diagram. Data sources; Kamiaso (Shimizu et al., 1996), However, TDMs from each district cluster around Oki-Dogo (Tanaka and Hoshino, 1987). Other data; Inomata (1999) and refer certain restricted ranges. It is noteworthy that the to caption of Fig. 6 in Lee et al. (2001). DM; Depleted Mantle, CHUR; TDMs (2.2 Ga) of southern Yamizo are completely Chondritic Uniform Reservoir. 148 H. Kagami et al. / Gondwana Research 9 (2006) 142-151 and/or 147Sm/144Nd>0.13 ratios), Mino belt (some sedimen- which indicates an inclination of Nd evolution line from 2.6 Ga to present, is a quite common for granitoids rocks (e.g., Table 1 tary rocks), Hida-Oki belt (some metasedimentary rocks) and in Kagami et al., 1992). Ashio belt (except for Yamizo district). Two samples from The Hida-Oki belt with TDMs of 2.6-2.45 Ga is probably Renge and Tamba belts also fall under the same age category. formed in and around the SKC as pointed out by Isozaki (1997) Granitoid rocks from Yanai and Kinki districts in the Ryoke and Ishiwatari and Tsujimori (2003). belt contain inherited zircon with ages of 1.95 Ga and 1.86 Ga (2) 2.3-2.05 Ga: Some sedimentary rocks from Kiso (Mino (Herzig et al., 1998). Jurassic mafic volcanic rocks occurring in belt) and from the southern Yamizo (Ashio belt) and some the Kinki Ryoke contain inherited zircon with an age of 1.9 Ga metasedimentary rocks from the Hida (Hida-Oki belt) belong (lizumi et al., unpublished data). Metasedimentary rocks from to this age category (Figs. 1,5). the Ryoke Kinki contain detrital zircon with a Pb-Pb age of ca. The major Proterozoic igneous activity in the SKC started 1.8 Ga (Ishizaka, 1969). These zircon ages are coincident with at ca. 2.1 Ga (e.g., Kim, 1987; Kim et al., 1995; Inomata, older (1.8 Ga~) TDMs from Ryoke belt. 1998). The oldest inherited zircon CHIME ages reported for Major Proterozoic igneous activities of the SKC are dated the gneisses from southern Korea Peninsula are 2.15 Ga around 2.0 and 1.7 Ga by Kim (1987) and 2.0-1.6 Ga by Kim (Suzuki et al., 1999). The Nd model ages from Yangtze et al. (1995). The rearranged TDMs for the YC is define an age Craton (abbreviated as YC), reported by Chen and Yang span of 2.0-1.6 Ga. The coincidence of TDMs among some 143Nd/144Nd<0.5125 and (2000), can be arranged as regions of the SW Japan Arc and the YC as well as SKC are 147Sm/144Nd<0.13 for calculation of model age, and the shown in Fig. 3, and we note a distinct overlap of Sm and Nd age values fall under three groups of 2.3-2.1, 2.0-1.6 and isotopic data from both regions. less than 1.5 Ga. Thus, it is noteworthy that the TDMs of The TDMs of age category of 1.9-1.55 Ga thus fall within 2.3-2.05 Ga are close to the beginning of the major igneous I 'K pue s u Anoe snout jo ueds su noeu a activity within the SKC and oldest group of the YC. Besides has been proposed by Kagami et al. (2000) that the Triassic to this, the sedimentary rocks with 2.3-2.05 Ga are plotted in Early Jurassic mafic volcanic rocks constituting the Ryoke belt the field of the SKC and are close to the YC in 143Nd/144Nd were formed from the continental lithospheric mantle. Further- vs. 147Sm/144Nd diagram (Fig. 3). more, the sedimentary rocks from the analyzed samples plot in (3) 1.9-1.55 Ga: The belts or districts within belts the fields of active and passive continental margins in the log belonging to this age category are as follows: Ryoke belt (K2O/NazO) vs. SiO2 (wt.%) discrimination diagram of Roser and Korsch (1986). We infer that the sediments within the Ryoke belt as well as some other belts that belong to the same 5 age category might have been sourced from terrains within or (a)Kamiaso oki around the SKC and/or YC. ! Four samples from the Hida metasedimentary rocks belong OL to age categories of (2) and (3) (Figs. 1,5). However, three 1.6 2.6 rocks from this belt plot between 0.5 and 1.0 Ga lines as shown (b）Southern Ashio,Southern Yamizo &Tsukuba in Fig. 3. These data suggest that source of some metasedi- NorthernAshio&NorthernYamizo mentary rocks are relatively young. Accordingly, the Hida metasediments were sourced formed from relatively young terrains ranging in age from Proterozoic to Early Paleozoic. 15 Ryokebelt,Minobelt Tamba,Renge,WesternSuo&EasternSuo 5. Provenances of sedimentary rocks with young TDMs Some sedimentary rocks collected from the districts belonging to 1.45-1.25 Ga (4) and 1.2-0.85 Ga (5) contain detrital zircons with older ages as described above. If all sedimentary rocks consist of mixtures of several different components with various ages as well as detrital zircons, the obtained young TDMs are meaningless. However, all sedi- mentary rocks belonging to age categories (4) and (5) have high ratios of 143Nd/144Nd (>0.5122, Table 1, Fig. 3) in spite 1.0 1.2 1.4 1.6 1.8 2.0 2.2 2.4 2.6 of normal 147Sm/144Nd ratios (0.10-0.13) displayed by the (c)Higo sedimentary rocks in the Inner Zone of SW Japan Arc. These isotopic data strongly suggest that their source materials formed at relatively young age. In other words, the precursors +1.6+ 1.8 +*2.0+2.2 +2.4* 2.6 1.0 1.2 1.4 did not have a prolonged pre-sedimentation history. Thus, even Age (Ga) if the sedimentary rocks were formed from source materials with varying ages, most the contribution to the sediments were Fig. 5. Frequencies of TDMs of the Inner Zone of Southwest Japan Arc. Data sources; refer to caption of Fig. 1. derived from rocks with younger ages. H. Kagami et al. / Gondwana Research 9 (2006) 142-151 149 The sedimentary rocks with young TDMs of 1.45-0.85 Ga considering its geological position against the Hida-Oki belt. (age categories; (4), (5)) are distributed in the margin of this The clastic materials constituting the sedimentary rocks (2.3- study area towards the southwestern region (Suo belt, Higo) of 2.05 Ga; Miso-gawa complex in the Mino belt, southern SW Japan and northeastern region (northern Yamizo) of NE Yamizo in the Ashio belt, Hida-Oki belt) are derived from the Japan except Sawando complex in Chubu district (Fig. 5). SKC and/or YC on the basis of the Sm and Nd isotopic (4) 1.45-1.25 Ga: The rocks in this age category are mainly compositions and the time spans of igneous activity. The clastic reported from the Suo belt. The TDMs of 1.45-1.25 Ga are materials constituting the sedimentary rocks with TDMs of coincident with ca. 1.4 Ga Neoproterozoic igneous activity 1.9-1.55 Ga of the Ryoke, Mino and Ashio belts are probably recorded from the SKC (Kim, 1987). However, the sedimen- derived from the SKC and/or YC. On the other hand, the rocks tary rocks with ages of categories of (4) as well as (5) plot with the young TDMS of 1.45-0.85 Ga (Suo belt, Higo, outside the fields of the SKC and YC as shown in Fig. 3. probably the Akiyoshi belt, Sawando complex in the Mino belt, One data (sample, no. 44) from the Tamba belt belongs to northern Yamizo in the Ashio belt) are mainly made up of the age category (4), but another one (no. 45) fall under clastic materials derived from the different blocks from the category (3). Similarly, one data (no. 1) from the Renge belt SKC and YC. Especially, the Higo is considered as a fragment belongs to category (3), and another taken from Owada et al. of collision zone, e.g., Dabie terrane, between two continents in (1999) relates to category (4). Thus, more TDMs data on these East China. two belts are needed to analyze their provenances. However, Source of the Ryoke Sedimentary Rocks with extremely the rocks from the Renge belt plot within or around the fields of SKC and YC as shown in Fig. 3. Judging from geological In this study, some of the Ryoke metasedimentary rocks PNt+i/wSti 1opue pNt/PNeti uu Ainxo ppioi position of this belt at the margin of the Hida-Oki belt (Fig. 1), its provenance is identified as the SKC. ratios (Table 1, Fig. 3). These rocks have anomalous chemical (5) 1.2-0.85 Ga: This age category is mainly defined by the composition as compared with other metasedimentary rocks rocks from the Higo (Fig. 5), which has been generally from the Ryoke belt, which are characterized by extremely high accepted as western extension of the Ryoke belt. However, as TiO2, total Fe2O3, MgO, P2O5, Cr and Ni, and low Zr/Ti and described above, the TDMs of the Higo are quite different from SiO2/AlzO3 ratios (Kawano et al., 2006). These chemical those of the typical Ryoke belt which implies that the former is characteristics together with the high 147Sm/144Nd ratios not a part of the latter as pointed out by Hara et al. (1992) and Osanai et al. (1996). igneous rocks. Mafic igneous rocks are found within the Ryoke Though the TDMs of 1.2-0.85 Ga overlap with one of the belt itself and also in Tamba belt. The former group was Neoproterozoic igneous activities in the SKC at 1.0-0.7 Ga emplaced during Triassic to Early Jurassic (Kagami et al., (Kim, 1987), the Sm and Nd isotopic characteristics of the 2000; Okano et al., 2000), and the latter during Late Paleozoic Higo metamorphic rocks are quite different from those of the (Sano and Tazaki, 1989). The metamafic igneous rocks, so- SKC as well as YC rocks (Fig. 3). Provenance of the Higo called metadiabase, are closely and widely distributed in the rocks is presumed to be from a crustal block of relatively southern area of the sampling sites from Mie Ryoke (Yoshi- younger age. Based on petrological and geochronological zawa et al., 1966). studies of high-grade metamorphic rocks occurring in this The Sm and Nd isotopic relation between the metasedimen- terrane, Osanai et al. (1999, 2006) proposed that the Higo is a PN++1/uSti pue PNtt /PNt1 uoy Aonxo um syoou K1e, fragment of the collision zone, e.g., Dabie terrane, in East ratios is shown in Fig. 6, where they plot within or in between China. Nishiyama et al. (2004) also proposed that the Higo is a the field of the Tamba or Ryoke mafic igneous rocks. These fragment of the Qinling-Dabie-Sulu collision zone. Though TDMs of rocks from the Akiyoshi belt could not be 0.5132 10[d sone1 PN+i/PNeti pue PNtti/uSvi u poureqo outside the fields of the SKC and YC and the latter ratios are slightly higher than those of the Higo in the same range of 30 147Sm/144Nd ratios (Fig. 3). These isotopic features of this belt 0.5128 30 suggest that its provenance has similar or slightly young ages compared with that of the Higo. 31 38 39 0.5124 口 口口口 6. Concluding remarks for the provenance of SW Japan 37 Arc Mafic igneous rocks 147Sm/ 144Nd from Ryoke belt The provenance of the rocks with TDMs of 2.6-1.8 Ga 0.5120 0.08 0.10 0.12 0.14 0.16 0.18 0.20 0.22 from Hida-Oki belt and Kamiaso conglomerates is probably the SKC as inferred on the basis of relationship between Fig. 6. Sm and Nd isotopic relationship between the metasedimentary rocks 147Sm/144Nd and 143Nd/144Nd ratios, Nd isotopic evolution of sno pe so p/Nt pue pN/ws u xa m rocks from the Ryoke and Tamba belts. Data sources; Ryoke belt (Kagami et the SKC and U-Pb zircon ages. The provenance of the Renge al., 2000), Tamba belt (Sano and Tazaki, 1989). Numbers with closed squares belt (TDM=1.83 Ga) must also be from within the SKC show number of samples in Table 1. 150 H. Kagami et al. / Gondwana Research 9 (2006) 142-151 relations further strengthen our inference that the source of the Inomata, M., 1999. Ages of the basement rocks of the northern part of pe pNnN Aonxa m s Ars the Korea Peninsula. Abst. 106th Annu. Meet. Geol. Soc. Japan, vol. 114 147Sm/144Nd ratios is mafic igneous rocks of the Tamba and (in Japanese). Ishiwatari, A., 1985a. Granulite-facies metacummulates of the Yakuno Ryoke belts. ophiolite; evidence for unusually thick oceanic crust. J. Petrol. 26, 1-30. Ishiwatari, A., 1985b. Igneous petrogenesis of the Yakuno ophiolite (Japan) Acknowledgements in the context of the diversity of ophiolites. Contrib. Mineral. Petrol. 89, 155-167. The authors would like to express their appreciation to Ishiwatari, A., Tsujimori, T., 2003. Paleozoic ophiolites and blueschists in Japan and Russian Primorye in the tectonic framework of East Asia: a Profs. K. Suzuki (Nagoya Univ.) and M. Inomata (Tokyo synthesis. Isl. Arc 12, 190-206. Univ. of Agriculture) for providing information of age data Ishizaka, K., 1969. U-Th-Pb ages of zircon from the Ryoke metamorphic and Drs. C. Nzolang (Nigata Univ.), K. Sajeev (Okayama terrain, Kinki District. J. Jpn. Assoc. Mineral., Petrol. Econ. Geol. 62, Univ. of Science), Profs. M. Pandit (Rajasthan Univ.) and Y. 191-197 (in Japanese with English abstract). Isozaki, Y., 1997. Contrasting two types of orogen in Permo-Triassic Japan; Osanai (Kyushu Univ.） for their critical and constructive accretionary versus collisional. Isl. Arc 6, 2-24. comments on the manuscript. We thank Profs. T. Nureki Isozaki, Y., Maruyama, S., 1991. Studies on orogeny based on plate tectonics in (Okayama Univ.), Y. Nishimura (Yamaguchi Univ.) and Mr. Japan and new geotectonic subdivision of the Japanese Island. J.Geogr. M. Kataura (Pacific Consultants Co. Ltd.) for providing 100, 697-761 (in Japanese with English abstract). Kagami, H, Yokose, H, Honma, H., 1989. 87Sr/86Sr and 143Nd/144Nd ratios samples from the Ryoke belt and western area of Chugoku of GSJ rock reference samples; JB-1a; JA-1 and JG-1a. Geochem. J. 23, district, and thank Prof. M. Santosh (Kochi Univ.) for 209-214. improving the manuscript. We also thank Prof. A. Ishiwatari Kagami, Iizumi, S., Tainosho, Y., Owada, M., 1992. Spatial variations of Sr and (Kanazawa Univ.) and an anonymous referee for critical Nd isotope ratios of Cretaceous-Paleogene granitoid rocks, Southwest reviews. This work was supported by a Grant-in-Aid for Japan Arc. Contrib. Mineral. Petrol. 112, 165-177. Scientific Research (No. 10440139) from the Ministry of Kagami, H., Yuhara, M., Iizumi, S., Tainosho, Y., Owada, M., Ikeda, Y., Okano, O., Ochi, S., Hayama, Y., Nureki, T., 2000. Continental basalts in Education, Sports, Science and Technology of Japan. the accretionary complexes of the Southwest Japan Arc; constraints from geochemical and Sr and Nd isotopic data of metadiabase. Isl. Arc 9, 3-20. References Kawano, Y, Akiyama, M., Ikawa, T., Roser, B.P., Imaoka, T., Ishioka, J., Yuhara, M., Hamamoto, T., Hayasaka, Y, Kagami, H., 2006. 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Quadrangle Series (1:50,000), Kanazawa (10) 62., Geol. Surv. Japan (in Japanese with English abstract). abstract). Tanaka, S., 1992. Origin and the Early Mesozoic granitic rocks in the Hida Osanai, Y., Hamamoto, T., Kagami, H., Suzuki, K., Owada, M., Kamei, A., Terrane, Japan, its implication for evolution of the continental crust. J. Sci. 1999. The Higo high-grade metamorphic rocks in Japan as a part of the Hiroshima Univ., Ser. C 9, 435-493. collisional terrane between the Sino-Korean and Yangtze Cratons. Tanaka, T., Hoshino, M., 1987. Sm-Nd ages of the Oki metamorphic rocks and Gondwana Res. 2, 599-601. their geological significance. Abst. 94th Annu. Meet. Geol. Soc. pp. 492 Osanai, Y., Owada, M., Kamei, A., Hamamoto, T., Kagami, H., Toyoshima, T., (in Japanese). Tran Ngoc Nam, 2006. The Higo metamorphic complex in Kyushu, Japan Tanaka, T., Togashi, S., Kamioka, H., Amakawa, H., Kagami, H., Hamamoto, as the fragment of PermoTriassic metamorphic complexes in East China. 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Neodymium isotopic compositions of Silurian Yakuno IN andesite. Two specimens from the Kohonomai metagabbros. J. Mineral. Petrol. Sci. 87, 272-282 (in Japanese with metamorphic rocks, the Oki Island, Southwest Japan. Geochem. J. 28, English abstract). 333-339. Sano, S., Tazaki, K., 1989. Greenstones in the Tamba belt. Mem. Geol. Soc. Yoshizawa, H., Nakajima, W., Ishizaka, K., 1966. The Ryoke metamorphic Jpn. 3, 53-67 (in Japanese with English abstract) Zone of the Kinki District, Southwest Japan accomplishment of a regional Sano, S., Hayasaka, Y., Tazaki, K., 2000. Geochemical characteristics of geological map. Mem. Coll. Sci., Univ. Kyoto, Ser. B 32, 437-454. Carboniferous greenstones in the Inner Zone of Southwest Japan. Isl. Arc 9, Yuhara, M., Kagami, H., Nagao, K., 2000. Geochronological characteriza- 81-96. tion and petrogenesis of granitoids in the Ryoke belt, Southwest Japan Shibata, K., Adachi, M., 1974. Rb-Sr whole-rock ages of Precambrian Arc; constraints from K-Ar, Rb-Sr and Sm-Nd systematics. Isl. Arc 9, metamorphic rocks in the Kamiaso conglomerate from Central Japan. Earth 64-80. Planet. Sci. Lett. 21, 277-287.
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Geothermics 102 (2022) 102379 Contents lists available at ScienceDirect GEOTHERMICS Geothermics ELSEVIER journalhomepage:wv ate/geothermics Gas geochemistry of geothermal fluids from the Hatchobaru geothermal field, Japan Jun-ichiro Ishibashi a,b,*, Kei Yamashita a 1, Keigo Kitamura , Yasuhiro Fujimitsu , Syogo Oshima d, Yumi Kiyota d b Ocean-Bottom Exploration Center, Kobe University, Kobe 658-0022, Japan Department ofEarthResources Engineering,Faculty of EngineeringKyushu University,Fukuoka 819-0395,Japan dGeothermal Department, West Japan Engineering Consultants, Inc.,Fukuoka 810-0004,Japan ARTICLEINFO ABSTRACT Keywords: We analyzed the chemical and isotopic composition of the geothermal fluids from the Hatchobaru geothermal Estimated reservoir temperature field in the Kyushu Island, Japan. Fluid chemistry showed similarity with that reported in earlier studies. Magmatic volatiles Chemical geothermometry provided an estimated reservoir temperature of 250-300 °C. Helium and carbon Carbon isotope exchange geothermometry isotopic ratios of the steam is likely to reflect the signature of the magmatic heat source. The apparent equi- librium temperature for carbon isotope exchange between CO2 and CH4 was 375-430 °C. The sulfur isotope data were in accordance with the idea that an acidic Cl-sO4-type fluid was modified from a neutral Cl-type fluid. 1. Introduction (Kiyota et al., 1996; Matsuda et al., 2000); however, the gas chemistry of geothermal fluids has rarely been reported. Geochemical study of geothermal fluids is a useful technique for We had an opportunity to obtain the geothermal fluid samples geothermal exploration. Chemical geothermometry is a conventional directly from the well heads of the Hatchobaru geothermal power plant, tool to estimate the temperature in the reservoir. It is based on the using a portable phase separator that can collect the vapor phase (steam) assumption that chemical interactions between the fluid and minerals of without loss as well as the liquid phase (hot water). In this paper, we the surrounding rocks attained or at least were close to equilibrium. report the results of the chemical (major elements and gas species) and Moreover, the isotopic composition of specific species can provide in- no n (d s formation on their source. Geochemical data often support a model Hatchobaru field. We also discuss the geochemical signature of the deep based on a geophysical survey, because they directly reflect the prop- region of the geothermal system from the obtained data. erties of the materials dissolved in the ascending fluid. Among them, gas species are considered to be mainly derived from the heat source 2. Geological background magma, and can provide insights into the deeper region of the crust than from the fluid reservoir. The Hatchobaru geothermal field is located on the northwestern The Hatchobaru geothermal field is located in the central part of flank of the Kuju Volcanoes in the central part of Kyushu Island, Japan Kyushu Island and hosts one of the largest geothermal power plants in (Fig. 1). The surrounding area is characterized by a graben structure, Japan. Geochemical studies have been conducted since the early stage of where active fault systems are concentrated. The heat source of the geothermal exploration in the 1960s (e.g., Hayashida and Ezima, 1970). Hatchobaru geothermal field has been associated with the late Quater- Previous studies revealed two types of geothermal fluids: a neutral nary volcanic activity of the Kuju Volcanoes. Volcanic activity began at Cl-type fluid from the main reservoir and a Cl-sO4-type fluid mainly about 2oo ka, and the most recent eruption was a small ash fall in recognized in the peripheral region (e.g., Shimada et al., 1985). Recent 1995-1996 (Kamata and Kobayashi, 1997). On the east of the Hatch- geochemical studies have focused on the acidic Cl-sO4-type fluid obaru geothermal field is Mt. Kurodake, which is a lava dome that * Corresponding author at: Ocean-Bottom Exploration Center, Kobe University, 5-1-1 Fumake-minamimachi, Higashinada, Kobe, 658-0022 Hyogo, Japan. E-mail address: ishibashi@port.kobe-u.ac.jp (J.-i. Ishibashi). 1 Present address: Kajima Technical Research Institute, Chofu 182-0036, Japan https://doi.org/10.1016/j.geothermics.2022.102379 Received 30 October 2021; Received in revised form 13 February 2022; Accepted 17 February 2022 Available online 10 March 2022 0375-6505/@ 2022 Elsevier Ltd. All rights reserved. J.-i. Ishibashi et al. Geothermics 102(2022)102379 formed at 1.6 ka. 3. Methods The Hatchobaru geothermal field is stratigraphically divided into four units: (1) Kuju volcanic rocks (Middle-Upper Pleistocene) consist- Geothermal fluid samples were collected from four production wells ing of hornblende-bearing andesite lava, (2) Hohi volcanic rocks (Lower in the Hatchobaru geothermal power plant. Sampling was conducted in Pleistocene) consisting of pyroxene-bearing andesite lava flows, (3) Usa October and November 2018. Steam (vapor phase) and hot water (liquid group (Miocene) consisting of altered andesitic lava flows and brecci- phase) samples were collected separately by introducing the geothermal ated tuff, and (4) basement rocks (pre-Tertiary) composed of granitic fluid discharge from each well head into a portable phase separator. and metamorphic rocks, in a descending order (Fujino and Yamasaki, The hot water samples were collected in a polypropylene bottle and 1985). filtered with a 0.45 μm membrane filter within one hour. The filtered The Hatchobaru geothermal system is mainly controlled by a frac- samples were divided into two bottles, one of which was acidified by HCl ture network consisting of a NW-SE trending and perpendicular NE-SW for cation analysis and the other for anion analysis. Chemical analysis of trending faults (Fig. 1). Previous studies have revealed that the main the hot water samples was conducted at Kyushu University. Major an- reservoir of geothermal fluid is located in the southeastern part of the ions (Cl and SO4) were analyzed using an ion chromatograph (DX-100, geothermal field, and the fluid ascends towards the northwest along the Thermo Scientific Dionex) after adequate dilution. Alkalinity analysis by fault system (e.g., Shimada et al., 1994). Based on this understanding, titration with 0.1 N HCl was conducted for the quantitative determi- production wells are arranged at depths of 1000-2300 m to collect the nation of HCO3. Major cations (Na, Mg, K, and Ca) together with other fluids for the power plant. dominant species (SiO2 and B) were analyzed using an ICP-OES 5100, Agilent Technology, Tokyo) after adequate dilution. The precision N.01 Mt. Sensu 130°E 40°E Mt. Kuroiwa Hatchobaru GPP Mt.Mimata Kuju-lwoyama Mt. Goto fumarolic area Mt. Ryoshi Mt.Hosshg Mt. Nakadake 0 1.0 2.0km d as u s e se us si an a (i nd rmd n a s ie a go d i i s faults and inferred faults are illustrated as solid and dashed lines with bars. The peaks of Quaternary volcanoes are shown as triangles. The shaded ellipse shows the locality of the fumaroles of the Kuju-Iwoyama Volcano from where fumarole samples were collected in the previous studies (Amita and Ohsawa, 2003; Saito et al., 2002; Marty et al., 1989). J.-i. Ishibashi et al. late Pliocene to Pleistocene sedimentary basins. Loading of crustal for major element analysis was approximately ± 10% based on repeated Similar to previous studies (e.g., Shimada et al., 1985; Kiyota et al., analysis of a standard material. The isotopic compositions of hydrogen 1996), two types of geothermal fluids were recognized in the Hatch- and oxygen were determined using a mass spectrometer following the obaru geothermal field, which could be clearly distinguished by the pH conventional method. The sulfur isotopic composition of SO4 was of the hot waters, even when measured at room temperature. The determined using an elemental analyzer/ mass spectrometer (EA/IRMS) neutral Cl-type fluid had a pH of 69, whereas the acidic Cl-SO4-type after precipitation of BaSO4. The precision for the isotope measurement fluid had a pH of 3, 4. The neutral Cl-type fluid was enriched exclusively was estimated as ± 1%o for hydrogen, ± 0.1%o for oxygen, and ± 0.2% in Cl, whereas the acidic Cl-SO4-type fluid contained some amount of for sulfur. SO (Fig. 2a). The cation compositions of the two types of geothermal The steam samples were collected following the conventional alkali fluids were distinguished from the Giggenbach diagram (Giggenbach, absorption protocol (Orsat method). After condensation of HO, the non- 1988). The composition of the neutral Cl-type fluid plotted in a region condensable gas was introduced into the Cd acetate solution to precip- that was close to the chemical equilibrium within the reservoir, whereas itate CdS, which was later used for the determination of HS content and the composition of the acidic Cl-SO-type fluid plotted in the region of sulfur isotope measurements. Further, the non-condensable gas was partial equilibrium (Fig. 2b). introduced into a KOH solution, and the CO2 content was determined The hydrogen and oxygen isotopic compositions of the studied from the difference in volume for the alkali absorption stage. The re- geothermal fluids are plotted in a 6D-818o diagram (Fig. 3). also shows sidual component (R-gas) was transferred into an evacuated bottle and the data of river water from the adjacent area (Amita and Ohsawa, 2003) provided for analysis using a gas chromatograph with a TCD (thermal tomography, Sato et al. (unpublished results) proposed that steeply to conductivity detector) to determine the composition of the minor gas revealed isotopic compositions that had clearly shifted from those of species (N2, O2, H2, CH4, Ar, and He). The helium isotopic composition meteoric water. However, the deviation was not as large as to overlap and He/Ne ratio was determined by samples collected separately into with the high-temperature fumaroles of the Kuju-Iwoyama Volcano specific bottles using a high resolution mass spectrometer after purifi- (shaded ellipse in Fig. 3). cation procedure. The carbon isotopic composition of CO2 was deter- The chemical geothermometer based on quartz solubility (Fournier, mined using a mass spectrometer with double inlets after precipitation 1977) for the four samples yielded reservoir temperatures in a range of BaCOg from the KOH solution.The carbon isotopic composition of from 255 to 270°C (Table 1).The estimated reservoir temperature from CH4 in the R-gas fraction was determined using a gas chromatograph/ the samples collected in 19801998 showed a wider range, but their combustion/mass spectrometer (GC/C/IRMS).The sulfur isotopic average value was in agreement with those from the 2018 samples. The composition of H2S was determined using an elemental analyzer/ mass reservoir temperature estimated by the Na-K-Ca geothermometer spectrometer (EA/IRMS) after conversion of the CdS into BaSO4. The (Fournier and Truesdell, 1973) showed a range from 260 to 300 °C, precision for gas species analysis was approximately ± 5%, with the which was in good agreement with the estimation by the quartz geo- exception of trace amounts. The precision for isotope measurement was thermometer. Yahara and Tokita (201o) had demonstrated that the of normal faults in Quaternary time has played a significant role in de- reservoir temperature of the Hatchobaru geothermal field has been stress-induced deflection of the lithosphere due to intraplate compres- 4. Results and discussion These fluid chemistry signatures are attributed to a stable geothermal system. Actually, chemical composition of the fluid from 4.1.Fluid chemistry Well 2H-21 did now show significant change between 1996 and 2018. Recharge of meteoric water would have been abundant, and the dis- Table 1 summarizes the analytical results of the hot water (liquid charging geothermal fluid would have continually experienced matured o n o ne pns) d upnou sods (as fluid-mineral interactions within the reservoir. The neutral Cl-type fluid element concentrations, and isotopic ratios. Table 1 also presents the hot could be denoted as the primary geothermal fluid, whereas the acidic Cl- water data of samples collected during 19801998 from a previous SO4-type fluid could be considered as the modified by an additional report (NEDO, 2000) for comparison. Hydrogen and oxygen isotopic sulfur-rich component based on the results of Giggenbach diagram. ratios were calculated to represent the single-phase geothermal fluid in Previous studies proposed dominant fluid-mineral interactions with the reservoir. The estimated reservoir temperatures by quartz solubility andesitic volcanic rocks, which overlie the granitic-metamorphic base- (Fournier, 1977) and by the Na-K-Ca geothermometer (Fournier and ment, based on a B/Cl molar ratio of ~0.05, and Sr isotopic ratios thrust wedges upon densified lower crust may promote rapid subsi- (Shimada et al., 1994). Table 1 Chemical and isotopic compositions of hot waters collected from the Hatchobaru power station. Well ID Sampling pH Na K Mg Ca C1 SO4 HCOs B Sea of Japan s180 Depth (km) INa-K- date (SO4) mafic rocks -10 (mg/ (mg/ Eleva /8u) /8u) (mg/ (%) (%) (%) (°C) (°C) Ca L) L) L) L) L) L) L) L) 2H-21 2018/10/18 7.7 1630 286 <d.1. 6.70 2620 80.0 Toyama 34.3 6.0 +18.4 260 30 H-32 2018/11/20 7.2 1630 251 12.3 2450 170 23.1 Moho 733 56 6.5 +18.3 255 261 H-33 2018/11/21 3.1 1050 179 87 6.89 1390 40 <d.l. 23.0 6S- 6.1 +23.4 261 266 H-34 2018/11/22 8.3 1190 258 <d.l. 3.61 tomography 80.0 82.6 879 8S- 6.6 +19.2 269 297 H-4* 1981/2/5 7.9 1995 226 2.0 83.8 3505 8.0 43.2 Uplift 56.6 4.7 +20.0 227 222 2H-21* 1996/3/1 6.4 1590 328 0.1 14.4 2750 85.8 15.0 NNW Toyama trough 53.9 6.7 +16.2 285 282 H-28* 1995/10/17 3.4 1293 260 15.0 16.2 1830 747 <d.l. 729 59.5 -7.2 +23.2 249 H-29 1997/9/11 3.2 1090 1360 228 12.7 Quaternary 1530 739 <d.1. 10 km No VE 637 -56.3 -7.0 +23.2 241 1980/8/25 276 HT-3 3.3 211 5.7 6.69 2230 MSL <d.1. 28.9 691 MSL n.d. n.d. 250 265 HT- 1982/4/20 8.6 1330 245 0.16 9.3 2240 71.0 227 32.5 1050 6.6 + 9.0 281 284 51* 2HD-1 1998/6/25 8.5 1070 234 0.01 5.59 1540 26.9 97.7 32.3 1120 59.6 5.5 +16.5 288 Uplift anomalies in the lower crust appear to be spatially correlated with the *Hydrogen and oxygen isotope composition was corrected to estimate isotope compositions of the reservoir. Data of the samples collected during 19801998 are cited from NEDO (2000). 3 J.-i. Ishibashi et al. Geothermics 102(2022) 102379 (a) SO4 (b) Na/1000 ●pH<5 pH>5 口 HCO3 K/100 VMg ℃ 8 9 20 Fig. 2. Major element compositions of the Hatchobaru geothermal fluids. (a) Anion composition in the trilinear diagram and (b) cation composition plotted in the Giggenbach diagram. Closed circles represent the acidic Cl-sO4-type fluid and open squares represent the neutral Cl-type fluid (smaller symbols indicate data of geothermal fluids collcted during 1980-1998 from NEDO (2000)). 4.2. Gas chemistry 18O -15 -10 Table 2 summarizes the analytical results of the steam (vapor phase) samples that include chemical and isotopic compositions of the gas species. The gas species composition was calculated in ppm volume, which is similar to the value of μmol/mol, from analytical results of the non-condensable gas or R-gas samples. The isotopic compositions of carbon for CO2 and CH4, sulfur for H2S, and helium are also listed in -40 Table 2. Composition of the gas species showed the typical signature of a Fumarole from water-dominated geothermal system. As shown in a H2O-CO2-H2S Kuju-lwoyama ternary diagram, the amount of non-condensable gas was less than a few 8 percent, and was mainly CO2 with minor H2S (Fig. 4a). Shimada et al. 09- (1985) had proposed that the CO2/H2S ratio of a fluid is associated with its pH, where the samples showed a high CO2/H2S ratio of ~15 for the neutral Cl-type fluid, whereas this ratio was lower (< 10) for the acidic Cl-SO4-type fluid. However, such an association was not distinctly G>Hd pH>5 observed in the studied samples (Fig. 4a). The relative composition of 80 the trace gas species is illustrated in a N2-Ar-He ternary diagram (Fig. 4b). In the diagram, the Hatchobaru geothermal fluid data are Fig. 3. Hydrogen and oxygen isotopic compositions of the Hatchobaru plotted along a trend between the air-saturated water and the magmatic geothermal fluids. Close circles and open square symbols are same as for Fig. 2. Cross marks indicate isotopic compositions of cold springs and river waters in component represented by the gas composition of the fumaroles of the the adjacent area, and the shaded ellipse shows the data range of high tem- Kuju-Iwoyama Volcano (shaded ellipse in Fig. 4b). pe e s e o omi-n sre aad Both the helium and carbon isotopic compositions of the geothermal fluids were similar to those of the magmatic component. The geothermal Ohsawa, 2003). fluids were characterized by an elevated helium isotopic composition (°He/*He = [7.1-8.7] × 10-6) with a substantially high He/Ne ratio (Fig. 5), which overlapped with the reported values from the fumaroles Table 2 Chemical and isotopic composition of steams collected from the Hatchobaru power station. Well Sampling H20 CO2 H2S N2 02 H2 CH4 Ar He 813c 813℃ 834s 3He/*He 4He/20Ne ID date (CO2) (CH4) (H2S) (vol (ppm) (ppm) Fig. 10. A fence diagram showing the relationship between the seismic tomography from Matsubara and Obara (2011) and the topography in the Hokuriku region. The color scale represents perturbations in the P-wave velocity. A map view of P-wave velocity perturbation at a depth of 20 km shows that the topographic domain boundary between the Hida (ppm) (ppm) (ppm) (udd) (%) (%) (%) %) 2H-21 2018/10/18 99.9 650 117 17.9 L0 2.2 1.0 0.3 0.006 -6.6 -25.3 -0.5 8.66 ± 18.3 0.08 H-32 2018/11/20 99.9 624 146 22.9 1.7 4.1 0.5 0.3 0.003 -6.7 -27.2 -3.7 18.8 0.08 H-33 2018/11/21 99.2 7140 351 100 6.7 8.6 35.8 1.3 0.086 -4.4 -23.5 -2.4 8.14 ± 54.1 0.09 H-34 2018/11/22 99.6 4010 202 54.6 1.4 5.7 18.9 0.8 0.038 -6.8 -24.5 -3.8 Mountains and theToyama trough is consistent with that between the low and high velocity anomalies. 62.5 0.07 J.-i. Ishibashi et al. Geothermics 102 (2022) 102379 (a) H2O /100 (b) N2 /100 ●pH<5 pH>5 Waterat300°C Air-saturated Waterat100°℃ Fumarole from Kuju-lwoyama Fumarole from Kuju-lwoyama CO2 10xH2S 10xHe Ar Fig. 4. Gas species composition of the Hatchobaru geothermal fluids. (a) Relationship between H2O, CO2 and H2S, and (b) relationship between N2, Ar, and He. Close circles and open square symbols are same as for Fig. 2 (smaller symbols indicate data of geothermal fluids collected during 1992-1998 from NEDO (2000). Cross marks in (b) indicate the value of the air-saturated water. The range of high temperature fumaroles from the Kuju-Iwoyama Volcano is shown in arrow marks in (a) and as shaded elipse in (b) (Data Sources: Amita and Ohsawa, 2003). s compositions of the studied samples, except for the sample from the 3He / He = 8.4x10-6 H-33 well (813C(CO2) = -4.4%o). The narrow range of isotopic ratios suggest limited opportunity for crustal contamination through the pro- cess of the expulsion of gas species from the heat source magma, involved into the geothermal fluid circulation, and transported to the 4He Air surface as dissolved species. Water He 4.3. Carbon and sulfur isotope systematics ●pH<5 pH>5 The carbon and sulfur isotopic compositions were determined for the paired gas species in this study. The isotopic data could provide tem- 0.1 d p d FELLL species undergo significant mutual interactions in the deep region. A 0.1 10 100 1000 4He / 20Ne geothermometer based on carbon isotope fractionation between CO2 and CH4 has been applied to several geothermal fields, assuming that Fig. 5. Relationship between ?He/*He and He/Ne for the Hatchobaru both species interact during the ascension of magmatic volatiles and/or geothermal fluids. Note that both axes are shown in a log scale. Close circles in the fluid reservoir (e.g., Fiebig et al., 2004). This geothermometer is and open square symbols are same as for Fig. 2. Mixing curve between the air- widely used as a carbon isotope measurement of trace amounts of CH4 saturated water (cross mark) and the endmember that has *He/*He = 8.4 × 10-6 is illustrated. A close triangle indicates data of high temperature fumaroles ever, the validity of this geothermometer is still questionable. Some from the Kuju-Iwoyama Volcano (Data Sources: Marty et al., 1989). recent studies argued a sluggish isotope exchange rate in natural systems due to a lack of efficient catalysts (e.g., McCollom, 2013). Others have of the Kuju-Iwoyama Volcano (Marty et al., 1989). These ratios were demonstrated an unpreventable contribution of thermal degradation of interpreted to be because of mixing between the magmatic component organic matter in a geothermal system (e.g., Fiebig et al., 2019). (?He/*He = 8.4 × 10-6) and the air-saturated water (Fig. 5). The carbon Although the results must be interpreted with caution, discussion about isotopic composition of CO2 was in the range of 813C(CO2) = -6.8 to the carbon isotopic composition of CH4 in the context of a geo- -4.4%o, which was consistent with the reported value of 813C(CO2) = thermometer could provide useful information about the deeper regions -8.0%o from the fumaroles of the Kuju-Iwoyama Volcano (Saito et al., of the geothermal system. 2002). The carbon isotopic data of CO2 and CH4 from the Hatchobaru The geochemical signatures of the gas species suggested that the -aa nn m i un nd a p o Hatchobaru geothermal system was affected by the subjacent magma rium temperatures (AETs) for carbon isotope exchange (Horita, 2001). with the contribution of magmatic volatiles (Hedenquist et al., 2018), The carbon isotope ratio of CH4 showed a narrow range (813C(CH4) = although this contribution was not evident in the SD-s18o diagram -27.2 to -23.5%o). The plots in Fig. 6 are aligned with the line parallel (Fig. 3), where the observed shift was attributed to the influence of hot to the apparent equilibrium, even though one acidic fluid sample (H-33 water that was injected into the adjacent subsurface through the rein- well) was scattered. The AET calculated from the isotope deviations of the four 2018 samples ranged from 375 to 430 °C, which is distinctly significant magmatic contribution (Fig. 4), and the elevated 3He/He higher than the estimated reservoir temperature of approximately ratio was comparable to the highest range recognized among hot springs 250-300 °C by chemical geothermometers. The carbon isotope sys- in the Kyushu Island which was attributed to the magma ascension into a tematics could be explained by interactions between CO2 and CH4 in a relatively shallow depth (Horiguchi and Matsuda, 2013). Notably, no deeper region beneath the fluid reservoir. Such deep interaction zone is e pe p ad rnu a uamq nns in accordance with a recent numerical model for the Hatchobaru ： J.-i Ishibashi et al. Geothermics 102 (2022) 102379 +20 S>Hd ● S>Hd● pH>5 pH>5 600℃ +10 400C. S 工 0 200°℃ S C 40 4 3 O1 1 10 -60 -20 -16 -12 -8 -10 0 +10 +20 +30 834S(SO4) (CO2) (%) Fig. 6. Relationship between 813C(CH4) and 813C(CO2) for the Hatchobaru Fig. 7. Relationship between 834s(H2S) and 834s(SO4) for the Hatchobaru geothermal fluids. Close circles and open square symbols are same as for Fig. 2. geothermal fluids. Close circles and open square symbols are same as for Fig. 2. n an a e s Dashed lines indicate apparent equilibrium temperature for carbon isotope exchange calculated from Horita (2001). exchange calculated from Ohmoto and Lasaga (1982). geothermal system, where fluid supply was projected to occur from a (2) The helium and carbon (and possibly sulfur) isotopic composi- deeper reservoir within the basement (Momita et al., 2000). tions as well as gas compositions were similar to the reported The sulfur isotope data for H2S and SO4 from the Hatchobaru fluid values from the fumaroles of the nascent Kuju-Iwoyama Volcano. are plotted in Fig. 7. Although sulfur isotope deviations are not often The gas chemistry strongly suggested that the geothermal system used in geothermometry, AETs for sulfur isotope exchange (Ohmoto and was substantially affected by the contribution of magmatic vol- Lasaga, 1982) were superimposed in Fig. 7. Notably, the neutral Cl-type atiles from the subjacent heat source. fluid and acidic Cl-sO4-type fluid showed similar sulfur isotope ratios of (3) Carbon isotope systematics yielded an apparent equilibrium H2S (834s(H2S) = -3.8 to -0.5%), which overlapped with the magmatic temperature of 375-430 °C, which is distinctly higher than the value. For the Cl-type fluid, the apparent equilibrium temperature estimated reservoir temperature. This result could be interpreted calculated from the isotope deviations was 265-320 °C, which was as reflecting the interaction between CO2 and CH4 in the deep comparable to the estimated reservoir temperature of approximately region beneath the reservoir. o s s a o H2S and SO4 during the process where the magmatic volatiles are the neutral Cl-type fluid and acidic Cl-sO4-type fluid. This dif- involved into the fluid reservoir, and might be the result of a dispro- ference is in accordance with the idea that the latter was modified portional reaction. The acidic Cl-sO4-type fluid showed not only a high by the contribution of an additional sulfur-rich component. +23.2 to +23.4%o) compared with that of the neutral Cl-type fluid (834s Funding (SO4) = +9.0 to +20.0%o). This sulfur isotope systematics would be in accordance with the idea that contribution of an additional sulfur-rich This study was conducted with support from the “Potential survey component such as low-temperature fluids in the shallow level is an and estimation of power generation of supercritical geothermal re- important factor for generation of the acid Cl-sO4 type fluid (Matsuda sources in East Japan and Kyushu, Japan (FY2018-2020)" project that et al., 2000). was funded by the New energy and industrial technology development organization (NEDO). 5. Summary CRediT authorship contribution statement We analyzed the liquid and vapor phases of the geothermal fluids obtained directly from the well heads of the Hatchobaru geothermal Jun-ichiro Ishibashi: Conceptualization, Validation, Writing - power plant. We conclude the following inferences from the fluid and original draft. Kei Yamashita: Investigation. Keigo Kitamura: gas chemistry of the studied samples from the Hatchobaru geothermal Conceptualization, Project administration. Yasuhiro Fujimitsu: Project system: administration, Funding acquisition. Syogo Oshima: Resources, Su- pervision. Yumi Kiyota: Supervision. (1) The major element composition and estimated reservoir tem- perature by conventional geothermometers showed little change Declaration of Competing Interest over the past few decades. This suggests the stability of the geothermal system where fluid-mineral interactions sufficiently The authors declare that they have no known competing financial matured close to the chemical equilibrium within the reservoir. interests or personal relationships that could have appeared to influence the work reported in this paper. J.-i. Ishibashi et al. Geothermics 102 (2022) 102379 Acknowledgment Horita, J., 2001. Carbon isotope exchange in the system CO2-CH4 at elevated temperatures. Geochim. Cosmochim. Acta 65, 1907-1919. Kamata, H., Kobayashi, T., 1997. The eruptive rate and history of Kuju Volcano in Japan We appreciate cooperation of the Kyushu Electric Power Co., Inc. in during the past 15,000 years. J. Volcanol. Geotherm. Res. 76, 163-171. ads an sxa a sp a und o d pe us Kiyota, Y., Matsuda, K., Shimada, K., 1996. Characterization of acid water in the Otake- thanks to Kyuden Sangyo Co., Inc. for their sampling operation and Hatchobaru geothermal field. In: Proceedings of the 17th Annual PNOC-EDC Geothermal Conference. Manila, Philippines. Energy Development Corporation, analysis of gas species. We thank anonymous reviewers for constructive Pp. 131-135. Manila. comments that helped improve the manuscript. Marty, B., Jambon, A., Sano, Y., 1989. Helium isotopes and CO2 in volcanic gases of Japan. Chem. Geol. 76, 25-40. Matsuda, K., Shimada, K., Kiyota, Y., 2000. Development of study methods for clarifying References formation mechanism and distribution of acid geothermal-fluid -case studies of geothermal areas in Kyushu-Japan. In: Proceedings of the World Geothermal Amita, K., Ohsawa, S., 2003. Mixing process of air and underground water into magmatic Congress. Kyushu Tohoku, Japan, pp. 1425-1430, 2000. gas discharged from Kuju-Iwoyama fumarolic area of Kuju Volcano, Central Kyushu, McCollom, T.M., 2013. Laboratory simulations of abiotic hydrocarbon formation in Japan. J. Geotherm. Res. Soc. Jpn. 25, 245-265 (in Japanese with English abstract). Earth's deep subsurface. Rev. Mineral. Geochem. 75, 467-494. Fiebig, J., Chiodini, G., Caliro, S., Rizzo, A., Spangenberg, J., Hunziker, J.C., 2004. Momita, M., Tokita, H., Matsuda, K., Takagi, H., Soeda, Y., Tosha, T., Koide, K., 2000. Chemical and isotope equilibrium between CO2 and CH4 in fumarolic gas discharges: Deep geothermal structure and the hydrothermal system in the Otake Hatchobaru generation of CH4 in arc magmatic-hydrothermal systems. Geochim. Cosmochim. geothermal field, Japan. In: Procedings of the 22nd New Zealand Geothermal Acta 68, 2221-2334. Workshop. Auckland, New Zealand, pp. 257-262. Fiebig, J., Stefansson, A., Ricci, A., Tassi, F., Viveiros, F., Silva, S., Lopez, T.M., NEDO, 2000. Report of Deep-Seated Geothermal Resources Survey Project. New Energy Schreiber, C., Hofmann, S., Mountain, B.W., 2019. Abiogenesis not required to Development Organization (NEDO), Kawasaki, Japan. explain the origin of volcanic-hydrothermal hydrocarbons. Geochem. Perspect. Lett. Ohmoto, H., Lasaga, A.C., 1982. Kinetics of reactions between aqueous sulfates and 11, 23-27. sulfides in hydrothermal systems. Geochim. Cosmochim. Acta 46, 1727-1745. Fournier, R.0., 1977. Chemical geothermometers and mixing models for geothermal Saito, G., Shinohara, H, Kazahaya, K., 2002. Successive sampling of fumarolic gases at systems. Geothermics 5, 41-50. Fournier, R.O., Truesdell, A.H., 1973. An empirical Na-K-Ca geothermometer for natural variations and precision of the gas sampling and analytical techniques. Geochem. J. waters. Geochim. Cosmochim. Acta 37, 1255-1275. 36,1-20. Fujino, T., Yamasaki, T., 1985. Geologic and geothermal structure of the Hatchobaru Shimada, K., Fujino, T., Koga, A., Hirowatari, K., 1985. Acid hot water discharging from geothermal wells in the Hatchobaru geothermal field. Chinetsu J. Jpn. Geotherm. Giggenbach, W.F., 1988. Geothermal solute equilibria: derivation of Na-K-Mg-Ca Energy Assoc. 22, 276-292 (in Japanese with English abstract). geoindicators. Geochim. Cosmochim. Acta 52, 2749-2765. Shimada, K., Tagomori, K., Fujino, T., Kitakoga, I., 1994. Deep geothermal structure and Hayashida, T., Ezima, Y., 1970. Development of Otake geothermal field. Geothermics 2, targets of future development in the Hatchobaru geothermal field. Chinetsu J. Jpn. 208-220. Hedenquist, J.W., Taguchi, S., Shinohara, H, 2018. Features of large magmatic- Yahara, T., Tokita, H., 2010. Sustainability of the Hatchobaru geothermal field, Japan. hydrothermal systems in Japan: characteristics similar to the tops of porphyry Geothermics 39, 382-390. copper deposits. Resour. Geol. 68, 164-180. Horiguchi, K., Matsuda, J., 2013. Geographical distribution of He/He ratios in north Kyushu, Japan: geophysical implications for the occurrence of mantle-derived fluids at deep crustal levels. Chem. 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Available online at www.sciencedirect.com SCIENCE Cdirect? LITHOS ELSEVIER Lithos 89 (2006) 47-65 www.elsevier.com/locate/lithos Geochemical and isotopic constraints on the genesis of the Permian ferropicritic rocks from the Mino-Tamba belt, SW Japan Yuji Ichiyama a*, Akira Ishiwatari ?, Yuka Hirahara b, Kenji Shuto c Department of Earth Sciences, Faculty of Science, Kanazawa University, Kakuma, Kanazawa, 920-1192, Japan b Graduate School of Science and Technology, Nigata University, 2-8050 Ikarashi, Nigata, 950-2181, Japan Department of Geology,Faculty of Science,Nigata University,2-8050 Ikarashi, Nigata,950-2181, Japan Received 14 February 2005; accepted 27 September 2005 Available online 14 November 2006 Abstract The Permian ferropicrite and picritic ferrobasalt occur in the Jurassic accretionary complexes of the Mino-Tamba belt as dikes intruded into the basaltic volcanic rocks. They are characterized by high MgO (11-27 wt.%), FeO* (16-20 wt.%) and HFSE rich nature is apparently magmatic in origin. The incompatible element contents and ratios indicate that the picritic ferrobasalt has degrees of partial melting of HFSE-enriched source material at high pressures (4-5 GPa). On the other hand, the ferropicrite may have been produced by the same degree of partial melting at a lower pressure, and subsequent olivine accumulation. The Sr and Nd isotopic signatures (*/Sr/86Sr(i)=0.70266 to 0.70329 and εNd(i)=+5.7 to +7.7) of these picritic and ferropicritic rocks are nearly constant and are equivalent to those of HIMU rocks, which require involvement of subducted oceanic crust material into their source region. Nevertheless, the ferropicritic melt cannot have been generated from the iron-poor picrite melt by crystal partial melting at extremely high pressures. A possible source material for the ferropicrite is the mixture of the recycled Fe- and Ti- massifs as Fe- and Ti-rich eclogite bodies. The ferropicritic magma may have been derived from the Permian, deep mantle plume in that the greenstone-limestone-chert complexes in the Mino-Tamba belt may be fragments of an oceanic plateau formed by the in the Jurassic time. 2005 Elsevier B.V. All rights reserved. Keywords: Ferropicrite; HIMU magma; Recycled oceanic crust; Permian mantle plume; Mino-Tamba belt; Sr and Nd isotopes 1. Introduction A ferropicrite is characterized by high-Mg and -Fe * Corresponding author. Present address: Society of Gem & Pre- (typically Fe0*>14 wt.%) content, and was first rec- cious Metal, Ohtsu 2094, Kofu, 400-0055, Japan. Tel.: +81 55 243 6147; fax: +81 55 243 6147. ognized in the Early Proterozoic Pechenga Group, NW E-mail address: yichi2san@juno.ocn.ne.jp (Y. Ichiyama). Russia by Hanski and Smolkin (1989). Francis et al. 0024-4937/$ - see front matter @ 2005 Elsevier B.V. Allrights reserved doi:10.1016/j.lithos.2005.09.006 48 Y.Ichiyamaetal./Lithos89(2006)47-65 (1999) suggested that the occurrences of the iron-rich picrite was derived from the Permian oceanic super- picritic and komatitic rocks reflect the high iron con- plume activity (see also a review by Ishiwatari and tent of the Precambrian mantle relative to the present Ichiyama, 2004). Recently, we have also found ferro- one. Recently, however, Phanerozoic ferropicrites are picritic rocks from the Mino-Tamba belt. These rocks discovered in some large igneous provinces (LIPs), and have unusually high Fe0* content (up to 20 wt.%), and it is suggested that the occurrence of ferropicrite is contain iron-rich mineral phases such as aegirine and restricted to the LIPs regardless of geologic age, and riebeckite. We believe that these rocks provide impor- that the ferropicrite is generated by partial melting of tant constraint not only on the origin of the Mino- mantle plume consisted of a basalt-peridotite mixture Tamba belt, but also on the Permian igneous activity (Gibson et al., 2000; Gibson, 2002). of the whole earth. In this paper, we report the petrol- An accretionary complex commonly consists of the ogy, mineralogy and geochemistry of the ferropicritic rock sequence composed of mafic volcanics (green- rocks, and discuss their petrogenetic relationship with stones), limestone, chert, sandstone and mudstone in the Mino HFSE-rich picrite, and other continental and younging order, forming the so-called “oceanic plate oceanic picrites and their geodynamic implications. stratigraphy", which were accreted to a continental mar- gin or an island arc by the long-term subduction of an 2. Geological outline and mode of occurrence oceanic plate. Therefore, the investigation of the accret- ed greenstone provides us indispensable information on Geologic structure of the Inner Zone of southwestern the ancient oceanic magmatism before Jurassic. The Japan is essentially constructed by nappe piles. The Mino-Tamba belt is one of the Jurassic accretionary Paleozoic-Mesozoic terranes of continental blocks complexes distributed in Japan (Fig. 1). Ichiyama and (Hida and Oki), ophiolites (Oeyama and Yakuno), Ishiwatari (2005) studied the Permian HFSE-rich picrite metamorphic belts (Unazuki, Renge, Suo and Ryoke) from the Mino belt, and discussed the genetic relation- and accretionary complexes (Akiyoshi, Maizuru, Ultra- ships with other terrestrial HFSE-rich picrites. They Tamba and Mino-Tamba) are tectonically overthrusted suggested that the HFSE-rich picrites were generated in order of increasing age as going upward in the nappe by the partial melting of a mantle plume material con- piles (Isozaki and Maruyama, 1991). Analogous struc- taining previously subducted and recycled oceanic crust ture is also seen in Primorye of Far East Russia, and it is fragments at polybaric pressures, and that the Mino interpreted that the Japanese and Russian nappe piles JapanSea Minobelt Fu MB Gifu Tamba Mino 8888 Shuzan Kumogahata Sakamoto-toge N Samondake Tambabelt Haiya Funafuseyama LakeBiwa Kuze Tsurugaoka Nabi Yuragawa Kanayama 10km Furuya F. Fig. 1. Distribution of the Jurassic accretionary complexes (gray) in East Asia and geotectonic map of the Mino-Tamba belt in the Inner Zone of southwestern Japan (Nakae, 2000). Ob and Fu indicate the Obama (Tamba) and Funafuseyama (Mino) areas, respectively. UTB and MB indicate the Ultra-Tamba and Maizuru belts, respectively. Y.Ichiyamaetal./Lithos89(2006)47-65 49 were detached from each other by the opening of the complexes are equivalent to Shuzan, Kumogahata, Japan Sea during Miocene (e.g. Ishiwatari and Tsuji- Haiya and Tsurugaoka complexes of the type area mori, 2003). (Fig. 1), respectively. The ferropicritic rocks occur in The Jurassic accretionary complexes are widely dis- the Kouchi (Haiya) complex. This complex is equiva- tributed in SW Japan (Fig. 1). It is divided into the Mino lent to the Funafuseyama complex in the Mino belt, and Tamba belts with the boundary at the Lake Biwa. where Ichiyama and Ishiwatari (2005) reported the The Mino-Tamba belt is tectonically overthrusted by HFSE-rich picrite and basanite sills and related hyalo- the Permian Ultra-Tamba belt, and is regionally meta- clastite interbedded withMiddlePermian chert at Funa- morphosed toward the Median Tectonic Line (Creta- fuseyama area (Fig. 1). The Kouchi complex consists of ceous low-P/T type Ryoke metamorphic belt). The the allochthonous slabs of mafic volcanics, chert, lime- Mino-Tamba belt also consists of some nappes, and is stone and siliceous mudstone and the melange com- subdivided into some tectonic units or complexes based posed of the blocks of these rocks in a mudstone on structure, lithology and age (e.g. Wakita, 1988; Iso- matrix. The structurally upper part of the complex is zaki, 1997; Nakae, 2000). The regional classification of dominated by melange, whereas the lower part is dom- complexes has been proposed for each of the Mino and inated by the large-scale slabs composed mainly of Tamba belts, but Nakae (2000) recently proposes a chert and basaltic rocks. The limestone, chert and mud- comprehensive correlation between the complexes in stone matrix of melange contain the Early to Middle both belts. We follow his proposal (Fig. 1). Permian fusulinids (Isomi and Kuroda, 1958; Sakagu- The ferropicritic rocks occur in the Obama area of chi et al., 1973), the Late Permian and Early to Middle theTamba belt.In the Obama area, Nakae and Jurassic radiolarians, and Middle Jurassic radiolarians Yoshioka (1998) subdivided the Tamba belt into some (Nakae and Yoshioka, 1998), respectively. Therefore, local complexes named Tada, Shimonegori, Kouchi and the age of the mafic volcanic rocks associated with both Mukugawa and the Late Jurassic Furuya Formation chert and limestone is Early to Middle Permian. The consisted of stratified clastic sequence (Fig. 2). These mafic rocks are basaltic massive and pillow lava, hya- Obama Tamba belt ShimonegoriComplex VITA Mixedrocks Mudstone Limestone Chert ☆ Picriticferrobasalt ★ Ferropicrite MukugawaComplex Thrust Hyaku e（931m） 1km KouchiC Fig. 2. Geological map of the Obama area in the Tamba belt (Nakae and Yoshioka, 1998) and sampling points of the studied ferropicrite and picritic ferrobasalt. 50 Y. Ichiyama et al. / Lithos 89 (2006) 47-65 loclastite and rare dikes, and are associated with chert Kanazawa University. The analytical procedure was beds and limestone blocks. The basaltic lava sometimes described by Shirasaka et al. (2004). The consistency includes chert xenoliths (up to 30 cm in diameter). Sano between INAA and ICP-MS methods is within about and Tazaki (1989) reported the Nd-Sm isochron ages of 20% for most elements. 303-339 Ma from the mafic volcanics and separated Sr and Nd isotopes of the rock and separated mineral clinopyroxenes in the Tamba belt. However, the error of samples were analyzed on a Finnigan MAT262 mass these ages ranges from ±40 to ± 88 Ma, and hence spectrometer at Department of Geology, Nigata Uni- these ages are of low reliability. In this paper, we versity. The analytical procedure follows Miyazaki and interpret the age of the mafic rocks as Early Permian Shuto (1998). The analyzed minerals (clinopyroxene (about 280 Ma) based on the fossil record of the and amphibole) were separated using isodynamic sepa- overlying chert and limestone as noted above. The ferropicrite frequently occurs as boulders in a 0.1194 and 146Nd/144Nd=0.7219, respectively. During stream of the Ojiga-tani valley, which is a branch of Onyugawa River (Fig. 2). In the upper stream of the the course of this study, the NBS987 Sr standard gave a (t=u) 2100000 560701L'0=I98/1S8 J0 9n1A ue Ojiga-tani, the boulders of the aegirine-bearing ankar- amite also occur. Judging from the distribution of these and the JNdi-1 Nd standard gave a mean value of boulders, it is likely that the ferropicrite and ankaramite 143Nd/144Nd=0.512060 ± 0.000013 (n=3). Sr, Rb, Nd occur as the intrusions (or lava) in basaltic rocks and/or and Sm contents were also determined by the same chert. The kaersutite-rich basalt also occurs as a dike PNost-uS6t1 pue IS+s-qd8 Suisn sadoos! se spouou intruding into the basaltic rocks. The picritic ferrobasalt mixed spikes. Sr and Rb contents of all whole rock occurs as a dike (about 40 m in thickness) associated samples are determined using XRF, and Nd and Sm with an aegirine-bearing gabbro dike intruded into the contents of some whole rock and separated minerals basaltic lava and hyaloclastite. samples are determined using ICP-MS. eNd was 3. Analytical techniques Mineral chemistry was analyzed by a wavelength- 4. Petrography and mineralogy dispersive EPMA (JEOL JXA-8800R) at the Coopera- tive Research Center of Kanazawa University. The 4.1. Ferropicrite analyses were performed under an accelerating voltage of 15 kV and a beam current of 15 nA. JEOL software Representative analyses of the minerals are shown in using ZAF corrections was employed for data reduc- Table 1. The ferropicrite includes an abundance of tion. Natural and synthetic minerals of known compo- olivine (40-45 vol.%) and clinopyroxene (9-11 sition are used as standards. The Fe3+ content in the Cr- vol.%) as phenocrysts. The olivine phenocrysts occur spinel was estimated by assuming spinel stoichiometry. as pseudomorphs completely replaced by secondary Major elements (Si, Ti, Al, Fe, Mn, Mg, Ca, Na, K carbonate, chlorite, talc and serpentinite. It is fine- and P) and trace elements (Ni, Sr, Rb, Ba, Nb, Zr, Y grained (commonly <1 mm in size) and anhedral to etc.) were analyzed by the Rigaku System 3270 X-ray euhedral. Some olivines are poikilitically enclosed by fluorescence spectrometer with Rh tube at Kanazawa clinopyroxene and amphibole (Fig. 3a). Opaque Cr- University for fused glass of Li2B4O7 (5 g)-rock pow- spinels are sometimes included in olivine, and are der (0.5 g) mixture and pressed pellet of rock powder (5 rimmed by magnetite. The Cr-spinels are 66-69 in g), respectively. The analyses were carried out at 50 kV Cr# (=100Cr/(Cr+A1),27-33 in Mg# (=100Mg/ accelerating voltage and 20 mA beam current. Rare (Mg+Fe2+)). and 27-30 in Fe3+# (=100Fe3+/ earth elements (REE) and some other trace elements (Cr+ Al+Fe3+). Some olivines contain spherical, com- (Cr, Ta, Hf, etc.） were determined by instrumental posite inclusions consisting of quenched augite, kaer- neutron activation analysis (INAA). The INAA samples sutite, biotite and magnesioriebeckite (rare) in a chlorite (0.1 g rock powder) were activated at the Kyoto Uni- matrix (altered glass) (Fig. 3b and c), which would have versity Reactor, and analyzed by a Ge (Li) -ray detec- been liquid inclusions trapped during the crystallization tor at the Radioisotope Laboratory for Natural Science of the host olivine. The clinopyroxene (<2 mm in size) and Technology, Kanazawa University. Two rock sam- shows the zoning from yellowish augite in core to ples were also analyzed by inductively coupled plasma pinkish titanaugite in rim. The clinopyroxene pheno- mass spectrometry using a Thermo Electron X7 at crysts are Mg#=82-86, TiO2=0.8-1.8 wt.% and Table 1 Representative mineral composition chemistry of the ferropicrite and picritic ferrobasalt from the Obama area, Tamba belt Ferropicrite Sample 040406-OG1 040522-0G8 Mineral Cpx core Rim Kaersutite Biotite Fe-rich augite Kaersutite Riebeckite Cpx core Rim Kaersutite Biotite Cr-spinelCr-spinel Kaersutite Cpx Phenocryst Cpx rim Groundmasss Olivine-inclusion Phenocryst Cpx rim Groundmasss Olivine-inclusion SiO2 50.78 46.43 44.46 37.69 44.15 39.71 52.36 52.16 51.64 41.66 38.69 0.06 0.01 39.71 50.63 TiO2 1.65 3.74 3.65 5.07 2.60 4.12 0.62 1.17 1.15 5.53 2.89 5.87 5.56 4.12 1.54 3.91 Al2O3 7.27 9.65 13.18 8.92 13.53 3.54 2.97 2.93 11.58 12.28 9.95 9.77 13.53 4.46 Cr2O3 0.41 0.00 0.00 0.01 0.03 0.00 0.95 0.13 1.00 0.00 0.04 28.73 30.10 0.03 0.00 FeO* 6.12 7.14 9.62 12.46 12.14 14.14 17.48 5.35 5.16 9.32 11.65 45.88 44.04 14.14 7.47 MnO 0.10 0.08 0.13 0.13 0.16 0.22 0.42 0.08 0.11 0.10 0.12 0.42 0.39 0.22 0.19 MgO 15.19 12.81 15.71 18.11 7.95 7.81 13.68 15.88 15.49 14.67 20.66 6.18 6.77 7.81 14.33 CaO 21.90 21.84 10.42 0.00 22.32 16.49 0.76 21.77 21.66 11.35 0.03 0.00 0.00 16.49 21.04 Na2O 0.45 0.58 3.68 0.11 0.56 1.94 8.79 0.43 0.42 3.20 0.09 0.00 0.00 1.94 0.66 Ichiy K20 0.00 0.00 0.62 8.26 0.00 0.33 0.14 0.01 0.00 0.57 7.46 0.00 0.00 0.33 0.00 98.8 98.31 93.88 96.65 yama Total 100.5 100.02 97.96 95 97.79 100.77 99.55 97.99 97.09 98.31 100.31 Mg# 0.82 0.76 0.81 0.72 0.54 0.52 0.95 0.84 0.84 0.76 0.76 0.28 0.31 0.52 0.77 Cr# 0.66 0.67 Fe3+# 0.30 0.29 / Lithos An% Picriticferrobasalt (2006) Sample 040523-T2 040523-7 Mineral Cpx core Rim Kaersutite Aegirine Biotite Hornblende Fe-richaugite Biotite Cpx core Cpx core Kaersutite Plagioclase Biotite Cpx Biotite Phenocryst Cpx rim Groundmasss Olivine-inclusion Phenocryst Cpx rim Phenocryst Groundmasss Olivine-inclusion 5 SiO2 49.32 46.27 40.68 53.37 36.89 42.25 52.42 37.31 47.40 49.61 39.97 57.19 35.37 45.70 33.73 TiO2 2.04 2.92 3.95 3.41 5.60 1.55 0.21 2.75 2.72 2.10 5.12 0.09 6.51 3.64 2.72 4.87 7.16 11.88 1.47 14.89 13.13 3.18 15.08 6.67 4.90 12.93 26.43 15.13 7.75 15.81 Cr2O3 0.02 0.01 0.00 0.03 0.01 0.00 0.01 0.00 0.01 0.02 0.04 0.03 0.01 0.01 0.00 FeO* 7.08 7.75 14.69 25.44 13.14 13.99 12.83 15.55 7.90 7.71 12.96 0.63 14.40 8.03 20.18 MnO 0.15 0.15 0.23 0.06 0.10 0.19 0.33 0.13 0.17 0.18 0.25 0.00 0.14 0.12 0.33 MgO 14.01 12.68 11.09 0.27 15.15 12.44 9.44 16.74 12.77 13.74 11.59 0.07 14.14 12.20 14.65 CaO 22.46 21.96 11.23 0.38 0.02 10.91 19.96 0.05 21.37 21.29 11.03 8.45 0.06 21.65 0.06 Na2O 0.41 0.45 2.95 14.06 0.79 2.81 1.95 0.72 0.46 2.90 6.01 0.86 0.48 0.48 0.28 0.02 0.02 0.05 8.40 0.85 0.03 7.48 0.00 1.12 0.51 7.79 K2O 1.04 0.01 0.00 4.96 Total 100.37 99.37 97.74 98.53 95.01 98.12 100.36 95.81 99.49 100.01 97.91 99.40 94.39 99.56 92.70 Mg# 0.78 0.74 0.59 0.67 0.74 0.57 0.66 0.74 0.76 0.62 0.64 0.73 0.56 Cr# Fe3+# An% 0.44 52 Y. Ichiyama et al. / Lithos 89 (2006) 47-65 >SE Fig.B in the liquid inclusions are characterized by a more Fe-, composed of altered glass (chlorite), plagioclase, apa- 150 tite and biotite. The abundant opaque minerals in the 4.7 wt.% and Al2Os 7.1-10.9 wt.%). The clinopyrox- groundmass (up to 11 vol.%) are Ti-magnetite and ene is rimmed by kaersutite, actinolite, cummingtonite ilmenite. Secondary titanite is also common. and rare magnesioriebeckite. Biotite, apatite and glass (chloritized) fill interstitial spaces. Ilmenite and Ti-mag- 4.3.Otherrocks netite are abundant (up to 4 vol.%). Secondary titanite is also common. The ankaramite includes large euhedral titanaugite phenocrysts (up to 3 mm in size; 45 vol.%). The 4.2.Picriticferrobasalt titanaugite (Mg#=69-80, TiO2=2.1-5.1 wt.% and Al2O3=5.4-9.6 wt.% in core) is rimmed by green The picritic ferrobasalt is rich in phenocrysts, which aegirine and aegirine-augite. The interstitial spaces are olivine (about 27-29 vol.%), titanaugite (20-22 between the large titanaugite phenocrysts are filled vol.%), plagioclase (16-23 vol.%) and rare kaersutite. with the groundmass composed of smaller grains of Olivine phenocrysts are pseudomorphs completely >tuff replaced by secondary chlorite and talc. They are fine- magnetite and ilmenite. The groundmass also includes grained (<1 mm) and subhedral or euhedral, and are altered glass replaced by chlorite and calcite. Minor often elongated. Some olivines contain the round-shaped biotite occurs as accessory minerals. Secondary titanite inclusions consisted of augite, kaersutite and biotite in a (A) chlorite matrix. The titanaugite (<1 mm; Mg#=73-78, (altered), kaersutite, clinopyroxene (altered) and plagio- TiO2=1.6-3.6 wt.% and Al2O3 =3.7-7.7 wt.% in core) is rimmed by kaersutite and rare aegirine (Fig. 3d). Most monly fine-grained (<1 mm), and kaersutite is the most plagioclase is replaced by sericite and chlorite, but rare abundant among them. Apatite and opaque mineral also relics are present. Plagioclase (An4o-s1) poikilitically occur as accessory minerals. The aegirine-bearing gab- a) b C a Fig. 3. Photomicrographs of the Tamba ferropicrite (a, b and c) and picritic ferrobasalt (d). All were taken with open nicol. Ol=olivine (pseudomorph), Cpx=clinopyroxene, Krs=kaersutite, Aeg=aegirine and Rbk=riebeckite. (a) Olivine poikilitically enclosed by clinopyroxene with kaersutite rim. (b) Needle-like clinopyroxene in the rounded “melt’ inclusion in olivine. (c) Riebeckite in the rounded “melt” inclusion in olivine. (b) Aegirine rimming clinopyroxene. Y. Ichiyama et al. / Lithos 89 (2006) 47-65 53 bro mainly consists of coarse-grained (around 3 mm in 5. Whole rock chemistry size) plagioclase, kaersutite and titanaugite (Mg#= 71- 76 TiO2= 1.4-2.9 wt.% and Al2O3 = 3.6-6.9 wt.%). The 5.1.Majorandtraceelements plagioclase is replaced by albite, K-feldspar, chlorite and calcite. Because titanaugite and kaersutite are in- The results of major and trace elements are listed in tensely replaced by chlorite, the estimate of the primary Table 2. Fig. 4 illustrates the major and trace element proportion of the two minerals is sometimes difficult. compositions with respect toMgO content for the Aegirine, aegirine-augite, apatite, biotite and Fe-Ti Tamba ferropicrite and picritic ferrobasalt, compared oxide occur as accessory minerals. Secondary titanite with those of other terrestrial picrites and ferropicrites. is also common. The elements mobile during alteration, such as K2O, Table 2 Whole rock chemistry of the ferropicrite and picritic ferrobasalt from the Obama area, Tamba belt. FeO* is total iron as FeO Rock Ferropicrite Picritic ferrobasalt Sample 040406-0G1 040522-0G8 040605-0G17 040605-OG19 040523-6 040523-7 040523-T2 SiO2 45.92 42.00 43.11 44.52 38.66 40.50 38.51 TiO2 1.18 1.25 1.30 1.31 4.71 4.34 4.72 Al2O3 5.03 4.60 4.96 5.06 10.12 12.62 10.32 FeO* 15.13 15.66 16.19 16.79 18.98 15.46 18.60 MnO 0.16 0.25 0.21 0.14 0.22 0.20 0.22 MgO 23.75 26.02 23.78 21.64 11.27 10.84 11.33 CaO 4.41 4.19 5.07 4.97 9.37 9.23 9.19 Na2O 0.20 0.39 0.34 0.35 the Pacific Basin. Jour. Geophys. Res., 89, 10291- 1.65 0.81 K2O 0.09 0.31 0.19 0.07 0.82 0.43 0.88 P2O5 0.32 0.32 0.31 0.34 0.68 0.88 0.69 Total 96.19 94.99 95.46 95.19 95.73 96.15 95.27 XRF XRF ICP-MS XRF XRF XRF XRF XRF ICP-MS Ni 1221 1182 1030 1133 1228 185 tinental margins. In “The Geology of Continental 199 174 Cu 36 25 30 49 30 38 35 Zn 138 lyzer. Jour. Geol. Soc. Japan, 82, 345-346, (in 126 176 185 1976: Pumpellyite-actinolite facies series of the 196 Pb 1 1.49 2 2 1.33 Rb 5 12 10.0 7 2 26 27 22.9 Sr 67 101 87.3 121 95 200 442 174 162 Y 17 16 12.1 17 19 31 36 29 24.7 Zr 103 112 99.2 1984: Relative motion between oceanic plates of 121 317 399 330 315 Nb 24 24 22.3 25 26 65 86 68 66.3 V 105 128 92.0 130 141 metabasic rocks from the Hast Schist Group of 438 575 366 Ba 75 121 57.2 74 86 463 433 406 284 INAA INAA INAA INAA INAA INAA INAA ICP-MS La 18.3 20.0 17.8 23.3 18.5 43.3 55.6 53.2 48.6 Ce 40.1 38.2 35.1 46.5 38.0 96.7 127 102 101 PN 17.2 21.1 57.4 69.6 58.3 48.5 Sm 3.49 3.81 3.67 4.36 3.88 9.33 11.8 9.60 9.42 Eu 1.40 1.34 1.23 1.20 1.31 2.61 3.86 3.26 2.94 Gd 3.74 8.82 Tb 0.515 1.13 Dy 2.78 5.85 Yb 1.38 1.58 1.00 1.68 1.64 2.13 3.12 2.34 2.09 Lu 0.142 0.139 0.183 0.216 0.276 0.294 0.265 0.293 Hf 2.31 2.31 2.38 2.98 2.29 6.68 8.27 7.07 7.41 Ta 1.32 1.55 1.58 1.20 1.64 4.54 6.32 4.47 5.19 Th 2.33 2.64 2.08 2.58 2.39 5.02 7.07 5.40 4.89 Cr 932 1219 914 1088 1444 419 289 421 311 Co 122 125 115 120 148 72.6 67.1 81.7 80.2 Sc 13.3 16.1 13.5 15.2 17.0 24.9 20.7 25.0 20.7 54 Y.Ichiyamaetal./Lithos89(2006)47-65 120 TiO2 (wt.%) Nb (ppm) 100 80 60 40 20 Al2O3 (wt.%)1 Zr (ppm) 4 12 10 300 200 100 Y (ppm) 35 12 30 10 25 20 15 10 5 CaO (wt.%) (wdd) !N) 1500 xx 1000 Xxx ++ 500 [FeO* (wt. %) 10 15 20 25 30 15 20 25 30 MgO (wt.%) MgO (wt.%) ●Tamba ferropicrite O Siberia Pechenga ■ Tamba picritic ferrobasalt East Greenland × Hawali ▲Mino picrite Etendeka + Iceland Fig. 4. Major and trace element variations with respect to MgO of the Tamba ferropicrite and picritic ferrobasalt. The plots are recalculated to anhydrous. For comparison, the picrites from the Mino belt (Ichiyama and Ishiwatari, 2005), Iceland (Skovgaard et al., 2001) and Hawaii (Norman and Garcia, 1999) and ferropicrites from Siberia (Arndt et al., 1995), Pechenga (Hanski and Smolkin, 1995), Etendeka (Gibson et al., 2000) and East Greenland (Peate et al., 2003) are also shown. Na2O, Ba, Pb and Sr, may have been changed during MgO (23-27 wt.%) and FeO* (16-18 wt.%) and low alteration processes, and in this paper, these elements TiO2, (1.2 wt.%), Al2O3 (5 wt.%) and CaO (5 wt.%) are not discussed in detail. In general, the major ele- contents. The picritic ferrobasalt shows lower MgO ments such as TiO2, AlzO3 and CaO of the Obama (%m o7 dn) *o q p (%m 71 puo) ferropicrite increase with the decrease of MgO content, than those of the ferropicrite. The TiO2 (4.5-5.0 wt.%), which is consistentwith accumulation or subtraction of Al2O3 (10-13 wt.%) and P2Os (0.7-0.9 wt.%) contents olivine. The Tamba ferropicrite is characterized by high are significantly high relative to those of the ferropicrite Y.Ichiyama et al. / Lithos 89 (2006) 4765 55 (TiO2=1.21.3wt.%,AlO=4.65.1wt.%and The contents of Nb (2426 ppm), Zr (103121 P2Os=0.3 wt.%).The FeO* content of the Tamba ppm) and Y (1619 ppm) in the ferropicrite slightly ferropicrite and picritic ferrobasalt is more than 14 increase with the decrease of MgO,but the Ni con- wt.%,and is classified to ferropicrite according to tent (11301230 ppm) does not show clear correlation Hanski (1992). This high FeO* content is clearly dif- with MgO. The picritic ferrobasalt is more enriched ferent from those of common picrites for example from in Nb (6586 ppm), Zr (317399 ppm) and Y (2936 Iceland and Hawaii.In addition,the comparison with ppm) and is poorer in Ni (132199 ppm) than those the other ferropicrites (Pechenga,Siberia and Etendeka) of the ferropicrite. The trace element patterns of the shows that the Tamba ferropicrite and picritic ferroba- ferropicrite and picritic ferrobasalt show the similar salt contain unusually high FeO* content. patterns enriched in more incompatible elements (Fig. 1000 (a) Tamba ferropicrite an Tamba picritic ferrobasalt man 100 Primitiv 10 S b 100 Mino Picrite Tamba Alkali d Sample Mino Tholeite Tamba Tholeite (C) O—Siberia Pechenga ——Etendeka +lceland mantl —EastGreeland —Hawaii 100 P 10 Id Sar 日 ThNbTaLa CeP H YYbLu Fig. 5. Primitive mantle-normalized trace element pattems of the Tamba ferropicrite and picritic ferrobasalt (a). For comparison, the patterms of (b) the Mino picrite (Ichiyama and Ishiwatari, 2005), MinoTamba basaltic rocks (Tamba tholeitic and alkali basalts (Sano et al., 2000) and Mino (Funafuseyama) tholeiitic basalts (Jones et al., 1993), and (c) the representative terrestrial picrites and ferropicrites (references are shown in the caption of Fig. 4) are also shown. Primitive mantle values are after Sun and McDonough (1989). 56 Y. Ichiyama et al. / Lithos 89 (2006) 47-65 Table 3 Sr and Nd isotopic composition for the picritic rocks from the Mino-Tamba belt 87Rb/ 87Sr/ 87Sr/ 147Sm/ 143Nd/ 143Nd/ Sample Rock Rb Sr 2 Sm PN 2 eNd 86Sr 86Sr 86Sr(i) 144Nd 144Nd(i) 144Ndi name (ppm) (ppm) (udd) (udd) 040522-0G8 Ferropicrite 12 101 0.344 0.704028 0.702732 3.76 17.6 0.129 0.512867 13 0.512644 6.77 CPX 0.363 114 0.009 0.702967 14 0.702932 6.43 23.4 0.166 0.512977 14 0.512689 7.66 CPX+AMP 1.82 275 0.019 0.703364 14 0.703291 10.9 46.4 0.142 0.512883 14 0.512637 6.63 040523-T2 Picritic ferrobasalt 27 174 0.449 0.704356 14 0.702664 9.42 48.5 0.117 0.512791 19 0.512587 5.67 CPX 1.98 132 0.043 0.702867 12 0.702703 9.88 37.6 0.159 0.512893 13 0.512617 6.26 CPX+AMP 12.8 235 0.157 0.703366 14 0.702774 14.5 69.2 0.127 0.512830 14 0.512610 6.12 021117-4L Mino picrite 7 195 0.104 0.703280 14 0.702888 8.52 39.4 0.131 0.512905 14 0.512678 7.44 CPX 8.94 29.9 0.181 0.512962 13 0.512648 6.86 CPX+AMP 14.6 4074 0.010 0.702769 0.702730 15.6 63.5 0.148 0.512915 11 0.512657 7.04 Italicized numbers represent the value determined by ICP-MS. Initial value is calculated at 265 Ma. 5). In more detail, the patterns of the picritic ferro- 5.2. Sr and Nd isotopes basalt show more depletion of Y and HREE than those of the ferropicrite, and have a slight positive Sr and Nd isotope and Rb, Sr, Nd and Sm contents anomaly of Ti. The contents of incompatible element are listed in Table 3. Fig. 6 shows the initial Sr isotopic ratio (87Sr/86Sra) and eNd?) plots for the whole rock such as Nb, Zr and TiO2 (HFSEs) in the Tamba ferropicrite are similar to those of the Hawaian and mineral separations of the ferropicrite and picritic picrite and the Pechenga and Etendeka ferropicrite, ferrobasalt and the Mino picrite. Initial isotopic ratios although the Nb content is somewhat higher. The are calculated on assumption of the 265 Ma age. This contents of these elements in the Tamba picritic age is based on the hyaloclastite and picrite associated ferrobasalt are very high and are similar to those of with the Middle Permian chert in the Funafuseyama the Mino picrite and the Siberian ferropicrite. ：The area, the Mino belt (Ichiyama and Ishiwatari, 2005). (143Nd/144Ndi) of these trace element patterns of the Tamba ferropicrite and The initial Nd isotopic ratios ( picritic ferrobasalt are very enriched in incompatible rocks and mineral separations show a nearly constant elements, and also are similar to that of the Mino value (0.51259-0.51269 and εNd(i) =+5.7 to +7.7). On picrite and Siberian ferropicrite. 15 Hydrothermal alteration * Tamba ferropicrite ■Tamba picritic ferrobasalt Mino picrite holeiite d Tamba alkali V 3 St.Helena HIMU EMII Siberia East Greenland BulkEarth EMI 5 0.702 0.703 0.704 0.705 0.706 0.707 0.708 Fig. 6. 87Sr/86Sr(@) vs. ENdi) composition of the Mino- Tamba picritic rocks. The compositional fields of representative volcanic rocks are shown as comparison (East Pacific MORB: White et al., 1987; Hawai: Stille et al., 1986; Tristan da Cunha: Le Roex et al., 1990; Samoa: Wright and White, 1987; Tubuai: Chauvel et al., 1992; St. Helena: Chaffey et al., 1989). The tholeitic and alkalic basalts are after Sano and Tazaki (1989). The ferropicrites from Siberia, East Greenland and Etendeka are from the same references as in Fig. 4. Y. Ichiyama et al. / Lithos 89 (2006) 47-65 57 some scattering (0.70266 to 0.70329). In particular, the d tectonic setting (Arai, 1992). In Fig. 7, the compositions range, which may have resulted from secondary alter- of Cr-spinel in the Tamba ferropicrite are shown with the ation. Because the 87Sr/86Sr(i) of the mineral separations available data of the terrestrial picrites and ferropicrites. of the ferropicrite must be nearer to the original value, The TiO2 content and Fe3+# ofthe Tamba ferropicrite are significantly higher than those of MORB (commonly<1 been high relative to the other rocks. As a whole, the Sr wt.% in TiO2 and Fe3+# <10), and indicate the affinity and Nd isotopic ratios show slightly depleted compo- with intra-plate basalts (Arai, 1992). These values are sition, and are plotted near the fields of MORB (DM) also clearly distinguished from those of Hawaii and Ice- and HIMU (high μ). land. The Cr-spinel of the Tamba ferropicrite is higher in Cr# and Fe3+# and lower in Mg# than those ofthe HFSE- 6. Discussion rich Mino picrite. The spinel compositions of the ferro- picrites from Pechenga, Etendeka, Siberia and Tamba 6.1. Iron-rich primary magma show wide scattering, but clearly indicate iron-rich char- acteristics such as higher Fe3+# and lower Mg# than To evaluate the MgO and FeO* contents in igneous those of Hawaii and Iceland picrites, which reflect the stage before alteration, the analyzed contents of the iron-rich nature of their host rocks. ferropicrite (040406-0G1) and 1picritic ferrobasalt Ishida et al. (1990) reported the iron-rich picrite (22 (040523-7) were compared with the result of simple wt.% in FeO*) highly sheared during the greenschist- mass balance calculation. The parameters used in this facies metamorphism, and thought that iron content of calculation are the approximate modal and chemical this rock may have been increased through metamor- compositions of the minerals in the two rocks (Table phism or later alteration. The mass balance calculation 4). The olivine compositions were estimated from the and the presence of iron-rich Cr-spinel and clinopyrox- coexisting clinopyroxene compositions, assuming that ene inclusions in olivine phenocrysts indicate that the the Mg-Fe exchange between olivine and clinopyrox- iron-rich nature of the Tamba ferropicritic rocks was ene is about 1:1 at high temperature (Loucks, 1996). inherited from iron-rich primary magma, not due to the The MgO content of glass was from the approximate enrichment of iron during secondary alteration. The melt compositions saturated in olivine + clinopyroxene ferropicrites underwent metamorphism lower than and olivine + clinopyroxene+ plagioclase, respectively, greenschist-facies grade, and contain aegirine and rie- which are identified by the experimental study at low beckite (Fig. 3). These minerals appear in high P/T pressures (Sack et al., 1987). The FeO* content of glass metamorphic rocks and differentiated alkaline igneous was estimated by assumption of the equilibration with rocks such as syenite. Although it is difficult to dis- olivine (Fe/Mg)°/(Fe/Mg)iq=0.3; Roeder and Emslie, criminate whether the aegirine and riebeckite are the 1970). The results (Table 4) are MgO=21.9 and products of either during igneous or metamorphic stage, FeO*=15.0 wt.% in the ferropicrite and MgO=14.2 their occurrence is consistent with the iron-rich char- and FeO*=14.8 wt.% in the picritic ferrobasalt, and acteristics of the ferropicrite. indicate the compositional consistency with analyzed MgO and FeO* contents within the range of the modal 6.2. Compositional differences among the ferropicrites error and compositional uncertainty. Cr-spinel is the first liquidus phase crystallizing from As shown in Figs. 4 and 5, there are compositional a cooling basaltic magma, and its composition is a good differences, especially in HFSE, among the terrestrial Table 4 Mass balance calculation for the ferropicrite and picritic ferrobasalt 10 Cpx P1 Kaer Bt bdo G1 Total Analyzed value Ferropicrite Mode (vol.%) 42 12 0 5 3 3 35 (040406-0G1) MgO (wt.%) 40.0 14.0 0.0 15.0 18.0 0.0 6.0 21.9 24.7 FeO* (wt.%) 20.0 7.0 0.0 10.0 10.0 50.0 10.0 15.0 15.7 Picritic ferrobasalt Mode (vol.%) 27 20 23 5 1 11 13 (040523-7) MgO (wt.%) 38.0 13.5 0.0 11.5 15.0 0.0 4.0 14.2 11.3 FeO* (wt.%) 22.0 8.0 0.0 13.0 14.0 50.0 7.5 14.8 16.1 58 Y.Ichiyamaetal./Lithos89(2006)47-65 12 Tamba MORB (eclogite)=2:1 ratio, respectively. In this dia- a ? 10 Mino gram, Zr content is changed by variable degrees of Siberia olivine accumulation or subtraction, while Nb/Zr ratio △ Etendeka (wt.9 Pechenga is not affected by this process. The ferropicrites are Hawaii divided into low Nb/Zr ratio (Etendeka and Pechenga) and high Nb/Zr ratio (Tamba, Siberia and East Green- 二 88 land) groups, suggesting different degrees of partial ? melting. If these ferropicrites contain no accumulated MORB olivine, the Tamba picritic ferrobasalt and the ferropi- 10 20 30 40 50 crites from Siberia and East Greenland cannot be Fe3 +/(Cr+Al+Fe3+ produced by the partial melting of primitive peridotite mantle due to their too high Zr content. Although it is 2 (b) highly dependant on the selection of the composition 10 and proportion of the source material, any enriched % component such as subducted oceanic crust may be (wt. necessary in the source region of these picritic mag- mas. Fig. 9 shows the correlation of Zr/Y and TiO2/ Al2O3 ratio, where both ratios increase with the in- creasing melting pressure (melt segregation depth) due 2 to decrease of AlzOs in clinopyroxene and the increas- MORB ing stability of garnet in residue with increasing pres- 20 40 60 80 100 sure (e.g. Walter, 1998). The Siberian and East Cr#（=100Cr/（Cr+AI)) Greenland ferropicrites clearly show higher Zr/Y and TiO2/AlzO3 ratios than those of Pechenga and Eten- +Fe2+y 100 Abyssalperidotite C deka, suggesting that the partial melting of the former 80 took place at greater depth relative to the latter. The g Tamba ferropicritic rocks show the ratios intermediate 100Mg/( 60 between these two groups, but the melting pressure Alpine-typeperidotite may have been closer to that of the Pechenga and 20 Etendeka ferropicrites. With respect to isotopic ratios (Fig. 6), the Etendeka IL 40 and East Greenland ferropicrite extend from HIMU to #6 M EMI component, while the Siberian ferropicrite are char- M 20 40 60 80 100 acterized by low 87sr/86Sr?) and intermediate eNd() Cr# (=100Cr/(Cr+AI)) which is similar to HIMU and the Mino-Tamba rocks, although the former is slightly higher in 87Sr/86Sr(i) and Fig. 7. Cr-spinel compositions of the Tamba ferropicrite and Mino picrite (Ichiyama and Ishiwatari, 2005; Ichiyama, unpublished data). lower in eNd(i) values than the Mino-Tamba rocks. Those of the picrites from Iceland (Sigurdsson et al., 2000) and Although not plotted, the eNd(i) value of+ 1.4 is reported Hawai (Nicholls and Stout, 1988) and the ferropicrites from Siberia for the Pechenga ferropicrite by Hanski et al. (1990). (Arndt et al., 1995), Etendeka (Gibson et al., 2000) and Pechenga (Hanski, 1992) are also shown. The fields of MORB are after Arai (1992). The compositional fields of the spinel in abyssal and alpine- 6.3.Relationship among theTamba ferropicrite,Tamba type peridotites are after Dick and Bullen (1984). picriticferrobasalt andMinopicrite In 87Sr/86Sr(i and eNda) diagram (Fig. 6),the Tamba ferropicrites. Although all ferropicrites are enriched in HFSE, the ferropicrites from Siberia, East Greenland ferropicritic rocks and Mino picrite are plotted in the and Tamba are significantly enriched in HFSE relative same space, suggesting that they are derived from the to those of Etendeka and Pechenga. Fig. 8 shows Nb/ source material with a similar composition. In addition, Zr ratio and Zr content of the Tamba ferropicritic these rocks show nearly constant Nb/Zr ratios (Fig. 8), rocks and the terrestrial ferropicrite. Two curves indi- cate the calculated compositional path of the two partial melting. Their different Zr contents may have melts produced by the batch melting of the primitive resulted from olivine fractionation, but they never share peridotite mantle and the mixed mantle at peridotite/ a common parent magma composition because any Y.Ichiyamaet al./Lithos 89(2006)47-65 59 0.4 Siberia 0.3 01 Pri EastGreenla 0.2 Olfra 0.1 ?5 Hawa ★★ Primitive mantle (2)+ Eclogte (1) Primitive mantie (2)+Fe-rich Eclogite (1) lceland partialmelts partial melks 100 200 300 400 500 Zr (ppm) Tamba ferropicrite ★ Primitive mantle ■ Tamba picritic ferrobasalt ★ Primitive mantle (2)+Eclogite (1) ▲ Mino picrite ★ Primitive mantle (2)+Fe-rich Eclogite (1) (references are shown in the caption of Fig. 4) are also shown. The paths of calculated partial melts for primitive mantle, primitive mantle +eclogite and primitive mantle+ Fe-rich eclogite are also shown. The partial melts are calculated by modal batch melting equation of Shaw (1970). Partition coefficients are after Halliday et al. (1995). The mineral proportion is olivine/orthopyroxene/clinopyroxene/garnet= 0.55:0.25:0.15:0.05 in primitive after Regelous et al. (1999). The small numbers near the paths indicate degree of partial melting. olivine control line cannot connect the ferropicrite and ratios of the samplewithout positive anomaly of Ti are picritic ferrobasalt (Fig. 4). The two picritic ferrobasalts similar to those of the Mino picrite. This leads to the are higher in TiO2/AlzOs ratio than the other in spite of context that theTambapicriticferrobasaltformed by the same Zr/Y ratio, which corresponds to the Ti pos- the melting of source material at the same pressure as itive anomalies in Fig. 5. The Zr/Y and TiO2/Al2O3 the Mino picrite, which is 4-5 GPa estimated by 100 Tamba ferropicrite Tamba picritic ferrobasalt Mino picrite East Greenland Pechenga. Siberia Hawaii Etendeka lceland 0.02 0.1 TiO2 Fig. 9. TiO2/AlzO3 vs. Zr/Y for the Tamba ferropicrite and picritic ferrobasalt. The representative terrestrial picrites and ferropicrites (references are shown in the caption of Fig. 4) are also shown. 60 Y.Ichiyamaetal./Lithos89(2006)47-65 Ichiyama and Ishiwatari (2005). In contrast to this, ferrobasalt occurs as poikilitic crystals including all these ratios of the ferropicrite are less than those of other mineral phases,indicating that it would have the Mino picrite. This is consistent with the absence of been the latest phase during crystallization. Therefore, the significant fractionation of HREE in the trace ele- we get into a dilemma that the Mino picrite shows the ment patterns of the ferropicrite, which indicates that same genetic condition as the Tamba picritic ferroba- the ferropicrite formed by the small degree of partial salt, but they cannot be related as a parent/daughter pair melting of the similar source material as for the picritic through crystal fractionation. ferrobasalt at shallower depth. It is likely that the elemental and isotopic relation- 6.4.Origin of the ferropicrite inthe Mino-Tamba belt ships between the Tamba picritic ferrobasalt and Mino The relatively low 87sr/86Sri) and high eNd(i) of the picrite indicate their strong genetic connection; that is, they were produced by the same degree of partial Tamba ferropicritic rocks and Mino picrite indicate the melting of a common source material at the same source composition similar to MORB or HIMU man- depth. Therefore, the more magnesian Mino picrite tle reservoirs, although a little EM components would can be a candidate of parental magma for the Tamba have been mixed with their source_ (particularly ferro- picritic ferrobasalt. To explain the daughter-parent re- - lationship between the picritic ferrobasalt and the Mino ering their high concentration of incompatible trace picrite, the crystal fractionation of olivine + plagioclase, elements and high Nb/Zr ratio, their source material is which is a crystal fractionation sequence observed in equivalent to the HIMU mantle source. The HIMU- MORB or some tholeitic layered complex (e.g. Mull), type volcanic rocks typically appear in Polynesia (e.g. might be the most reasonable mechanism to abruptly Tubuai Island) and St. Helena (Chauvel et al., 1992; increase FeO* content in the magma (Fig. 10). How- Chaffey et al., 1989), and it has been proposed that its ever, there is no such possibility of simultaneous crys- source material is the ancient subducted oceanic lith- tallizationof olivine and plagioclase, because osphere (e.g. Zindler and Hart, 1986; Weaver, 1991; plagioclase is very rare in the Mino picrite (Ichiyama Chauvel et al., 1992). Among the available isotopic and Ishiwatari, 2005), and plagioclase in the picritic data of the terrestrial ferropicrites, the Siberian ferro- Tamba terropicrite Tambapicriticferrobasalt Minopicrite 20 MORB+Peridotite(1:1) melt(3.5GPa) helt(5.0GPa) S wt. KLB-1 10 KG2 KG1 1.5GPa 1.5GPa Fog0 1GPa 5 10 15 20 25 30 MgO (wt.%) and HIMU-type picrites (Kogiso et al., 1997) are also shown. The compositions of melt from KLB-1 (circle) are after Hirose and Kushiro (1993) and Walter (1998). Those from KG1 (KLB-1/MORB=2:1; triangle) and KG2 (KLB-1/MORB=1:1; square) are after Kogiso et al. (1998). The MORB (GA1) and MORB pyrolite (MPY90) by Yaxley and Green (1998). The open star shows the melt composition (at 1650 °C and 5.0 GPa) from garnet pyroxenite by Kogiso et al. (2003). The arrows indicate the lines tying olivine with variable composition and the origin. Y. Ichiyama et al. / Lithos 89 (2006) 47-65 61 picrite is plotted near HIMU fields, and the other partial melting of such source material can no longer ferropicrites extend into EMI-EMII direction along produce iron-rich magnesian melt. Terrestrial ferropi- low εNd?) side of mantle array. This indicates that crites more or less contain igneous hydrous phases such their sources consist of the variable proportion of the as phlogopite and kaersutite as phenocryst and/or mixture of HIMU and EM, and involved pelagic groundmass minerals. This means that the source ma- sediments in addition to subducted oceanic crust terial of ferropicrite commonly contains HzO. However, (e.g. Dostal et al., 1998). In the previous studies for the melts from the melting experiment of hydrous ferropicrite genesis, Walker et al. (1997) examined peridotite show lower FeO* content than those of dry Re-Os isotope of the Proterozoic ferropicrites from peridotite (Kawamoto and Holloway, 1997). It is un- Pechenga, and revealed that the ferropicrite was orig- likely that the iron-rich nature of ferropicrite is due to inated from the plume mantle source with large pro- melting of a wet mantle material (Gibson, 2002). The portions of recycled oceanic crust. Gibson (2002) recycling of ferromanganese crusts and nodules with suggested that the iron-rich picritic melt is produced oceanic crust into mantle (Baker and Jensen, 2004) can by the partial melting of basalt+peridotite mixture at produce the iron-rich picritic melt. Nevertheless, much ≥1450 °C and ≥4.5 GPa. Arndt et al.(1998) and amount of ferromanganese material not only make the Arndt (2003） also interpreted that the meimechite picritic melt richer in iron but also make it highly associated with the Siberian ferropicrite was produced manganiferous and very high in Sr isotopic ratio. The bythe low degree (2%) of partial melting of latter features have never been observed among the MORB+peridotite at great depth (>6-7 GPa). ferropicrites. As mentioned above, the unusual iron-rich nature of The most reasonable explanation for the genesis of the Tamba ferropicrite and picritic ferrobasalt cannot be ferropicrite is the melting of the recycled iron- (and explained by crystal fractionation from a parental titanium-) rich basalt and/or gabbro. It is known that magma with an ordinary iron content. In the melting the occurrence of iron-rich basalt, gabbro and glass experiments of primitive (pyrolitic) mantle, the iron (up to 20 wt.% in FeO*) are widespread in ocean floor content of partial melts is sensitive to the change of such as East Pacific Rise, Indian Ridge and Galapagos pressure rather than the degree of partial melting, and Ridge (e.g. Regelous et al., 1999; Embley et al., 1988; increases with increasing pressure (e.g. Hirose and Coogan et al., 2001). In addition, the iron-rich (~20 Kushiro, 1993; Walter, 1998). The melt compositions wt.% in FeO*) eclogite with evolved MORB che- obtained by the melting experiments of the various mistry also occurs in some peridotite massifs (e.g. materials are shown in Fig. 10. The melt compositions the Voltri massif, Alps: Mottana and Bocchio, 1975; in this diagram are of the lowest degree of partial Ernst et al., 1983), and this indicates that iron-rich melting at each melting pressure (<16%), because the MORBs were actually subducted to the deeper mantle. ferropicrites used in this study appear to have been The involvement of the iron-rich MORB can produce produced by relatively low degrees of melting (Fig. the unusual high-FeO* picritic melt without significant 8). If the mantle source was pyrolitic, the Tamba ferro- modification of HIMU isotopic characteristics. Fig. picrite and picritic ferrobasalt require surprisingly deep 8 shows the compositional path of the melts from melt segregation depth (>7 GPa), which is inconsistent the mixture of primitive mantle+Fe-rich eclogite (Fe- with the estimated pressure in Zr/Y-TiO2/AlzOs dia- rich basalt). The path is placed to the right of another gram (Fig. 9). The HIMU-type basalt trends to be more path for a mixture of peridotite+eclogite (MORB), enriched in iron than other OIBs (e.g. Kogiso et al., which indicate that the melts produced by the melting 1997), and it is likely that the involvement of subducted of peridotite+Fe-rich eclogite (basalt) are more oceanic crust in the mantle source region can produce enriched in Zr at a given degree of partial melting. iron-rich partial melt. This is confirmed by the melting The picritic ferrobasalts are plotted between the two experiment of the mixture of MORB + peridotite, which compositional paths, fitting to our modeling. If the results in higher iron content than that of primitive recycled oceanic crust is Fe-rich gabbroic rock that mantle (Kogiso et al., 1998). Therefore, the involve- is poorer in trace elements, the compositional path ment of subducted oceanic crust component may play a should have shifted more to the left. The Fe-rich very important role in the genesis of ferropicrite. How- gabbros from the ocean floor show the positive Ti ever, the unusually high FeO* content up to 20 wt.% of anomaly (e.g. Coogan et al., 2001), and the positive the Tamba picritic ferrobasalt requires more MORB anomaly of Ti observed in the picritic ferrobasalt and component than 50% on the basis of the melting ex- some other ferropicrites (Fig. 5) would have been periment of the mixture of MORB+peridotite, but the inherited from the Fe, Ti-rich gabbro. 62 Y. Ichiyama et al. / Lithos 89 (2006) 47-65 The genesis of the Tamba ferropicrite is ambiguous, formed after the main basaltic volcanic activities. The because its high MgO content requires the partial melt- ferropicrites from Siberia and East Greenland also ap- ing at a high pressure (possibly up to 7 GPa; Fig. 10) pear as dikes and lavas later than the eruption of even if the subducted oceanic crust was involved in its voluminous continental flood basalt (Arndt et al., source region. This is inconsistent with the shallower 1995; Peate et al., 2003). These ferropicrites also melt segregation depth than that of the picritic ferroba- have compositional characteristics indicating the low salt and Mino picrite, because these rocks and the degrees of partial melting (Fig. 8). This suggests that a ferropicrite show the same Nb/Zr and initial isotopic colder mantle plume tail with recycled oceanic crust ratio. The high MgO content of the ferropicrite would can melt to produce ferropicrite after the arrival of a have resulted from olivine accumulation. This is consis- hotter mantle plume head. tent with the petrographic observations that the ferropi- It has been suggested that most mafic volcanic rocks crite shows phyric and poikilitic texture, and contains in the Mino-Tamba accretionary complex represent abundant olivine and clinopyroxene. In this case, if a remnants of many seamounts or oceanic islands in the half of the olivine phenocrysts in the ferropicrite (20 paleo-Pacific ocean in Carboniferous to Permian. Sano vol.%) is accumulated crystals, assuming average F078 and Tazaki (1989) and Sano et al. (2000) studied the composition (approximately MgO=40 wt.% and greenstones from the Tamba belt, and concluded that FeO=20 wt.%), the increase of 4 wt.% in FeO* results the greenstones were derived from the ocean islands from the accumulation. However, the size of the olivine composed of tholeitic and alkali basalt and from the phenocryst in the ferropicrite is commonly <1 mm, and ocean floor composed of N-MORB. Jones et al. (1993) it is unlikely that the accumulation of a large amount of suggested that the tholeitic volcanic rocks composi- olivine took place in the ferropicrite. Additionally, the tionally equivalent to E-MORB from the Funafuseyama possible accumulation of olivine + clinopyroxene causes area, Mino belt were originated from axis-centered less important increase FeO* than pure olivine accumu- oceanic plateau with a bathymetrically elevated feature lation. Therefore, it seems that the parental magma of like Iceland on the basis of their chemical composition the ferropicrite would also have contained higher FeO*, and association with huge limestone reefs. and that the recycling of iron-rich MORB into mantle It is difficult to decipher the original volcanic edi- source region should be required. fices of the greenstones from their fragmental occur- Although peridotites, pyroxenites and peridotite/ rences among the accretionary complexes. However, eclogite (MORB) mixtures have ever been experimen- the occurrence of relatively rare volcanic rocks such tally molten to determine the composition of partial as ferropicrite may help us to understand the origin of melts, any of the starting materials, except for the the greenstones and the ancient oceanic magmatism. Martian mantle enriched in iron (Bertka and Holloway, The plume-related ocean island basalt is mainly alkalic 1994), cannot produce the melts with the composition and is associated with fractionated rocks such as pho- comparable to terrestrial ferropicrites, especially the nolite and trachyte, while the greenstones from the Tamba ferropicrite. Our model is essentially consistent Mino-Tamba belt are dominated by tholeitic to alka- with the previous ideas that the source mantle of ferro- line basalt, and the fractionated rocks are very rare. In picrite may contain subducted ancient oceanic crust, but most LIPs, tholeiitic to alkalic basalts are also domi- we prefer the recycling of the Fe-rich basalt with nant. Ontong Java plateau mainly consists of tholeiitic evolved MORB composition to produce the ferropicri- basalt, but HIMU-like ferromagnesian basalts tic melt. (MgO=11 wt.% and Fe0*=14 wt.% in anhydrous) are also reported (Tejada et al., 1996). Gibson (2002) 6.5.Geodynamic implication of the ferropicrite reviewed the occurrence of terrestrial ferropicrites from Archean to Phanerozoic, and indicated that they typi- The geological, petrological and isotopic evidences cally occur in LIPs. Ichiyama and Ishiwatari (2005) clearly indicate that the Tamba ferropicritic rocks were pointed out that HFSE-rich picrites, including the derived from the Permian oceanic mantle plume. Fer- Mino picrite, typically occur in superplume region, ropicrites are rare in Phanerozoic. Gibson (2002) em- and that the presence of the Mino HFSE-rich picrite phasized that ferropicrite occurs at basal part of a thick indicates that the basaltic rocks in the Mino belt would volcanic sequence, and it is a consequence of the be produced by the upwelling of Permian superplume. melting of the margin of a mantle plume head. How- The occurrence of the ferropicrite and picritic ferroba- ever, the Tamba ferropicritic rocks clearly intruded into salt of the Permian age in the Jurassic Mino-Tamba more voluminous basaltic rocks, suggesting that they accretionary complex indicates a possibility that the Y. Ichiyama et al. / Lithos 89 (2006) 47-65 63 Mino-Tamba accretionary complex includes many (4) The Tamba ferropicritic rocks were derived from fragments of a Permian LIP that accreted into the the Permian mantle plume in an oceanic setting, subduction zone in the Jurassic time. The ancient oce- and may have formed after the arrival of its plume anic plateaux accreted to continental margin or arc have head. The plural occurrences of the ferropicritic been recognized in some accretionary complexes rocks and HFSE-rich picrite from the Mino- (Wrangellia terrane, North America: Lassiter et al., Tamba accretionary complex imply that this ac- 1995; Sorachi-Yezo belt, Japan: Kimura et al., 1994; cretionary complex would be fragments of an Tatsumi et al., 1998; Nagahashi and Miyashita, 2002), oceanic plateau formed by the Permian super- and parts of the present-day Ontong Java and Caribbean plume activities on paleo-Pacific ocean, and was plateaux indeed are obducted to Solomon Islands (e.g. subsequently accreted to the continental margin Tejada et al., 1996) and South America (e.g. Kerr et al., by subduction in the Jurassic time. 1997), respectively. If so, the greenstones of the Mino- Tamba accretionary complex would have been an oce- Acknowledgements anic plateau formed in paleo-Pacific ocean, which may be a new Late Paleozoic LIP other than the two known We thank S. Arai and T. Morishita of Kanazawa Permian continental flood basalt provinces (Siberia, University for several discussions. We are grateful for Russia, and Emeishan, China). More detailed petrolog- constructive reviews by two anonymous reviewers. We ical and geochemical studies are required to understand thank J. Takada and K. Takamiya of the Kyoto Univer- geodynamics of the Permian oceans. sity Research Reactor Institute and T. Nakanishi, Y. Nagamura, K. Washiyama and R. Amano of Kanazawa 7. Conclusion University for their help for INAA analysis. K. Koi- zumi is thanked for the assistance in the fieldwork. M. (1) The Permian ferropicritic rocks occur in the Shirasaka and Y. Shimizu also are thanked for the Mino-Tamba belt as dikes intruded into the assistance of ICP-MS and EPMA analysis, respective- mafic rocks. They are characterized by high ly. We also acknowledge that the Master thesis of T. MgO, FeO* and HFSE contents, and the miner- Muto (2001) of Kanazawa University inspired our re- alogical evidence indicates that their unusual search in the Obama area. iron-rich nature is of magmatic origin. (2) The incompatible element contents and ratios References indicate that the picritic ferrobasalt has strong genetic kinship with the highly HFSE-rich Mino Arai, S., 1992. Chemistry of chromian spinel in volcanic rocks as a picrite, and they were produced by the low degree potential guide to magma chemistry. Mineral. Mag. 56, 173-184. of partial melting of HFSE-enriched source ma- Arndt, N., 2003. Komatite, kimberlite, and boninite. J. Geophys. Res. terial at a high pressure. On the other hand, the 108B6, 2293. Arndt, N., Lehnert, K., Vasil'ev, Y., 1995. Meimechites: highly mag- Tamba ferropicrite was produced by the same nesian alkaline magmas from the subcontinental lithosphere? degree of partial melting at a lower pressure, Lithos 34, 41-59. and would be affected by olivine accumulation. Arndt, N., Chauvel, C., Czamanske, G., Fedorenko, V., 1998. Two The Sr and Nd isotopic signatures of the Tamba mantle sources, two plumbing systems: tholeitic and alkaline ferropicritic rocks and the Mino picrite are nearly magmatism of the Maymecha River basin, Siberian flood volcanic province. Contrib. Mineral. Petrol. 133, 297-313. homogeneous and indicate HIMU source. Never- Baker, J.A., Jensen, K.K., 2004. Cupled 1860s_187Os enrichments theless, the Tamba ferropicritic rocks cannot be in the Earth's mantle-core-mantle interaction or recycling of fractionated from the Mino picrite. ferromanganese crusts and nodules? Earth Planet. Sci. Lett. (3) To produce the unusually iron-rich picritic melt, 220,277-286. unreasonable addition of basaltic component or Bertka, C.M., Holloway, J.R., 1994. Anhydrous partial melting of an iron-rich mantle: II. Primary melt compositions at 15 kbar. Con- partial melting at unexpectedly high pressure is trib. Mineral. Petrol. 115, 323-338. required. The most reasonable source material of Chaffey, D.J., Cliff, R.A., Wilson, B.M., 1989. Characterization of the the ferropicrite is the mixture of recycled Fe- and St. Helena magma source. In: Saunders, A.D., Norry, M.J. (Eds.), Ti-rich basalt (and/or gabbro) and mantle perido- Magmatism in the Ocean Basins. Geol. Soc. Spec. Publ. vol. 42, pp. 257-276. tite. Such a ferrobasalt and its metamorphosed Chauvel, C., Hofmann, A.W., Vidal, P., 1992. HIMU-EM: the French equivalent occur in the present ocean floor and Polynesian connection. Earth Planet. Sci. 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Available online at www.sciencedirect.com TECTONOPHYSICS SCIENC DIRECT ELSEVIER Tectonophysics 395 (2005) 81-97 www.elsevier.com/locate/tecto Deformation mode in the frontal edge of an arc-arc collision zone: subsurface geology, active faults and paleomagnetism in southern central Hokkaido, Japan Yasuto Itoha*, Tatsuya Ishiyamab, Yasuhiko Nagasakic aDepartment of Earth Sciences,College of Integrated Arts andSciences,Osaka PrefectureUniversity, Gakuen-cho 1-1, Sakai, Osaka 599-8531, Japan AIST-logiclSuvey ofapan,Cenal 7thHgashi-1,Tsub 305567,aan TechnicalDepartment,Japan National Oil Corporation,Uchisaiwai-cho 2-2-2,Chiyoda-ku,Tokyo100-8511,Japan Received 10 September 2003; accepted 6 September 2004 Abstract Frontal edge of a Cenozoic arc-arc collision zone in southern central Hokkaido is described based on subsurface geology. Reflection seismic and drilling surveys delineate geologic structure of the N-S fold-and-thrust belt. West-verging faulted blocks extend as long as 200 km in central Hokkaido, and the activity of thrust has been propagating westward since the late Neogene. downthrown block of a floor thrust of the collision front and paleo-northeast Japan arc. Remanent magnetization of core samples belonging to the upthrown block indicates post-Oligocene counterclockwise rotation. It is opposite to surface rotational motions of the same block, then more detail study on fault architecture and block rotation is required to understand the mode of deformation of the fold-and-thrust belt. Maturity level of organic matters in a borehole suggests that burial by the thrusting initiated in the Quaternary. Thus, the deformation front of the arc-arc collision is in an active, nascent stage of crustal contraction. ① 2004 Elsevier B.V. All rights reserved. Keywords: Hokkaido; Arc-arc collision; Seismic interpretation; Borehole; Active fault; Paleomagnetism 1. Introduction of arc-arc collision events since the early Tertiary. The older collision between paleo-North American Southern central Hokkaido, northernmost compo- and Eurasian Plates initiated in the Eocene and nent of the Japanese Islands (Fig. 1) has been a site resulted in metamorphism under the Hidaka Moun- tains (Komatsu et al., 1983, 1989). A right-lateral * Corresponding author. Fax: +81 72 254 9752. component of motion is postulated for the contrac- E-mail address: itoh@el.cias.osakafu-u.ac.jp (Y. Itoh). tional event based on the fabric of metamorphic 0040-1951/S - see front matter @ 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2004.09.003 82 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 120° 130° 140° 150° LEGEND (a) (a) KurilBasin M moving direction Eurasian offore-arcsliver Hokkaido Plate (Fig.1b) directionofplate subduction convergent plate JapanSea boundary Pacific Plate Philippine 500km Sea Plate (b) LEGEND(b) oshi assumedblockof rotation(Takeuchi et al.,1999) terraneboundary Quaternary Neogene pre-Neogene 50km Zone Fig. 1. Maps showing tectonic context around the Japanese Islands (a) and geologic belts in Hokkaido (b; after Kato et al., 1990). rocks (Arita et al., 1986). The younger collision was of the constituents of Hokkaido. Takeuchi et al. caused by the westward migration of fore-arc sliver (1999) regarded central Hokkaido as blocks of of the Kuril arc (Fig. 1) since the late Miocene clockwise rotation based on paleomagnetic research (Kimura and Tamaki, 1985), which enhanced uplift (Fig. 1b). Their hypothetical blocks, however, have of the Hidaka Mountains and resulted in crustal not been verified through geologic structures indi- shortening of ~60 km (Kazuka et al., 2002). cative of shearing boundary or coherency of intra- Between such events, Kuril and Japan back-arc block paleomagnetic directions. basins began to spread (Lallemand and Jolivet, Thus, a series of tectonic events along the eastern 1986) inevitably, raising rotation and rearrangement Eurasian margin and the formation process of Y.Itoh et al./ Tectonophysics 395(2005) 81-97 83 14130 Ishikari Bay (Japan Sea) UmaoiHills Legend Alluvium Pleistocenesediments Umaoi Fm.(Pliocene) ueFm.(lo usrocks Oligocene-Eoce hingeofactive folds faults anticline X syncline solidcircles)andwithinOligocene (solid circes) 5 10km Pacific Ocean Fig. 2. Geologic map around the Umaoi anticline redrawn from Geological Survey of Japan (2002). Location of active fault and/or fold scarps (after Ikeda et al., 2002) are also shown. “u" and “d” attached on fault traces are upthrown and downthrown sides of faults, respectively. Sampling points of surface paleomagnetic data is after Kodama et al. (1993). 84 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 continental crust are recorded in the complex active thrust, Ishikari lowland on the west of collision geologic features in central Hokkaido. In this article, zone is covered by thick sedimentation. In an attempt the authors attempt to elucidate deformation mode to find hydrocarbon reserves, numerous seismic and within the western margin of the collision zone, drilling surveys have been organized in the area around the Umaoi Hills (Fig. 2), where active (Ishiwada et al., 1992). Important drilling sites and reverse faulting is observed (Ikeda et al., 2002). It seismic lines interpreted in this article are shown in is interpreted as the nascent portion of the westward Fig. 2. propagating crustal contraction (Ito, 2000). From We adopted two boreholes for geologic and another point of view, our study area belongs to the paleomagnetic analyses. MITI Umaoi (UM) was clockwise-rotated Block I (Fig. 1b) after Takeuchi et drilled on the Umaoi Hills, and penetrated thrust al.(1999). Based on subsurface information from sheets of the Paleogene units. In this study, seismic and drilling surveys, we delineate geologic paleomagnetic results for the Paleogene Minami- structures beneath the collision front. Then, paleo- naganuma Formation and the igneous basement magnetic analyses are executed on core samples samples beneath the floor thrust will be presented obtained from boreholes in order to assess rotational in the following sections. Eniwa SK-1 (EN) was motions around the fold-and-thrust belt. Because drilled in the Ishikari lowland and reached the core recovery in case of exploration is quite low earliest Cretaceous volcanic/sedimentary complex (generally less than 2%) unlike academic drilling (Suzuki et al., 1999). We collected samples from (e.g., ODP), we have made an effort to obtain the Cretaceous sedimentary rocks, of which direc- samples as much as possible. Notwithstanding low tional markers (bedding plane, fracture) are available recovery rate of core samples, deep exploration to determine core orientation and attitude of tectonic drilling under mud circulation with physical proper- tilt. ties continuously obtained through logging provides us with a unique information to clarify subsurface 2.2.Active faults structure. Our data require differential rotations within From a standpoint of tectonic geomorphology, the collision zone (or the hypothetical coherent the Umaoi Hills are a geomorphic expression of block), suggesting complicated modes of deforma- active folding at the leading edge of the fold-and- tion. Finally, degrees of organic maturation of thrust belt along the western margin of the Hidaka borehole samples are simulated using a kinetic Collision Zone. Recent deformations of the Umaoi model. The result gives us an estimate for the Hills are recorded in flights of fluvial and marine duration of thrust activity on the active collision front. isotope stage of 5 and 7 (Koike and Machida, 2001) that cover western flank of an anticline (Fig. 2; Ikeda et al., 2002). Numerous flexural 2. Background slip faults that cut and deform the terraces also suggest that a strong contraction occurs within 2.1. Geology forelimb of the anticline beneath the hills. Pyroclastic flow deposits at 40-42 ka,which Fig. 2 summarizes geologic units (Kato et al., mantle the southern Umaoi Hills, are also incor- 1990; Geological Survey of Japan, 2002) and faults porated in structural growth of the anticline (Ikeda around the study area. Umaoi Hills exist along et al., 2002). In addition, fold scarps and active western margin of the Hidaka Collision Zone. It folds upon the depositional surfaces of the consists of faulted anticlines of the Neogene rocks pyroclastic flow are continuously traced about 5 with narrow (under mappable size) exposures of the km to the west of the southern Umaoi Hills Oligocene Minami-naganuma Formation (Kurita and (Ikeda et al., 2002), indicating that pairs of Yokoi, 2000), which thickly develops beneath the imbricate thrusting occur at the leading edge Umaoi Hills and Ishikari lowland. Divided by an during the late Pleistocene time. Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 85 3. Seismic interpretation subsurface geologic structure (Kurita and Yokoi, 2000; Ishiwada et al., 1992; Taketomi and Nishita, 3.1.Umaoi section 2002). In 1995, a seismic survey was conducted around the coastal area by Ministry of Economy, Fig. 3 shows a multichannel E-W seismic profile Trade and Industry, (METI) Japan (JNOC, 1996). across the Umaoi Hills. In 1995, a seismic survey was During the shooting, 80 channels of hydrophones conducted by JAPEX. During the shooting, 330 (with an interval of 25 m) recorded the energy channels of geophones (with an interval of 25 m) released from a 24.6 1 (1500 in3) tuned airgun array, recorded the energy released from six vibrators, shot shot at 25-m interval. Raw seismic data were stacked at 50-m interval. Raw seismic data were stacked and and then subjected to a post-stack processing sequence in order to enhance the resolution. The SI order to enhance the resolution. 95-A seismic profile transects the southward projec- MITI Umaoi and some exploration boreholes have tion of pairs of active folds around the Umaoi Hills. been drilled upon or near this seismic line, and their Interpretation of the section (JNOC, 1996) shows geologic data (JNOC, 1997) provide important con- that the seismic data successfully images pairs of straints on an interpretation. As stated by Kurita and west-verging anticlines. These folds are underlain by a Yokoi (2000), the N-S trending hills are intensively west-verging thrust with shallow dips that crosscuts folded and faulted with westward vergence. Thrust the early Miocene strata (Takinoue Formation). Uni- blocks and their internal geologic structure were form thickness of the Karumai Formation in late confirmed as repeated Paleogene sections and attitude Miocene throughout the seismic section, identified by of bedding planes within the MITI Umaoi (Kurita and continuous reflectors, indicates that the Miocene strata Yokoi, 2000; JNOC, 1997), respectively. The floor predate the essential structural growth of the folds and thrust coincides with a boundary between the tilted are involved in hangingwall blocks. thrust horses and flat-lying horst-graben of the Tabular beds of pregrowth (late Miocene-Pliocene) Paleogene and Cretaceous igneous basement. formations are overlain by the Quaternary strata that It is noteworthy that the Neogene/Quaternary on taper hindward and onlap onto forelimbs and back- western flank of the anticlinal form are clearly limbs. These geometric relationships suggest that the deformed, suggesting thrust activity until recent Quaternary has buried structural relief of the under- period. Ages of such growth strata provide a clue to lying formations, and has been folded and rotated reconstruct history of fault motion and tectonic events, during their deposition. Hence, growth (Quaternary) thus will be examined in Discussion section. On the strata juxtaposed with either forelimbs or backlimbs other hand, the middle Miocene Kawabata Formation indicate structural growth of active folds during the on the eastern flank is involved into fault-bend fold, Quaternary. and truncated by the Karumai Formation in the late Miocene. Although the east-dipping fault in the 3.3.Active thrusting related to arc-arc collision eastern part of the section cuts the Karumai For- mation, the flat-lying strata imply that fault activity in The present seismic interpretation demonstrates the eastern area declined in late Neogene. Therefore, that a N-S thrust has been activated through the late thrusts bounding the Hidaka Collision Zone seem to Cenozoic bounding the western margin of the Hidaka propagate westward through the long-standing colli- Collision Zone. To the north, active landforms of the sion event as suggested by Ito (2000). thrust can be traced as far as 43°30'N (Ikeda et al., 2002). Southern extension of the thrust is identified as 3.2. Southern offshore section the deformation front of offshore fold zone (Ito, 2000) as far as 42°N. Therefore, the tectonic boundary is at Fig. 4 shows a multichannel E-W seismic profile least 200-km long, and comparable with width of the along the southern coast of central Hokkaido where Kuril fore-arc sliver (Kimura, 1996) colliding against economic reserves of hydrocarbon were confirmed the continental lithosphere of Hokkaido. A series of and numerous boreholes have been drilled to clarify microearthquake observations (Katsumata et al., WEST EAST MITIUmaoi 1km 200 400 600 800 1000 1200 1400CDPNo. 0 time(sec.) WO- Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 KarumaiFm. Nina Fm.+Quaternary (sec. waytime M.M. Upper Neogene L.-M.Miocene U.Oligocene(Minami- Cretaceous andQuaternary fault (TakinoueFm.) naganumaFm.) igneous basement Fig. 3. E-W multichannel seismic profile 95V-1 (migrated; upper) and its geologic interpretation (lower) running near the MITI Umaoi (at CDP no.550). Subsurface fault architecture and unit boundaries are based on interpretation of the exploration drilling (JNOC, 1997; Kurita and Yokoi, 2000). Paleomagnetic sampling horizons (UM2, 6, 7) are shown on the projected trace of the MITI Umaoi after depth to time conversion. Interval of maturity modeling (MM) is also shown on the same profile. See Fig. 2 for seismic line location. (a) CMPNumbers East 600 1500 1400 1300 1200 1100 1000 900 800 700 600 00 0.0 2.0 0 0 TT (sec) Y.Itohet al./Tectonophysics395(2005)81-97 .0 (b) CMPNumbers East 100 000 900 0.C 0 .0 W 5.0 6.0 2km Fig. 4. E-W multichannel seismic profile SI95-A (migrated; upper) and its geologic interpretation (lower) along the coast of southern central Hokkaido. Subsurface unit boundaries are after JNOC (1996). See Fig. 2 for seismic line location. 88 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 2002) showed that thrust-type focal mechanism was images. The Oligocene core samples (UM2-1~6) of dominant around our study area. In order to assess the the MITI Umaoi show rather wide ranges of dip future development of the arc-arc collision, deforma- angles (20~44°), which is comparable with those of tion history of the active and nascent fold-and-thrust logging data (28~47°). Because the dipping directions belt is of great importance. around the sampling interval are quite constant (289~291°), cores can be oriented with high reli- ability. As for the Cretaceous igneous rocks (UM6, 7) 4. Paleomagnetism of the MITI Umaoi well, samples were not initially oriented because sidewall imaging data were not Core samples for paleomagnetic analyses were available. Cylindrical rock specimens for magnetic obtained from the Eniwa SK-1 and MITI Umaoi measurements of 25 mm in diameter and 22 mm in boreholes, both of which had been drilled nearly length were cut from the core samples avoiding areas vertically. Although core recovery intervals of these with dense fracture or serious alteration. wells aimed at hydrocarbon exploration were quite Bulk magnetic susceptibility was measured for all short, subsurface samples with structural information specimens using a Bartington susceptibility meter provide us with an opportunity to study three-dimen- (MS2). Natural remanent magnetizations (NRM) of sional rotation scheme. Sampling depth, lithology and all specimens were measured with a 2-G Enterprise geologic ages (Kurita and Yokoi, 2000; Suzuki et al., three-axis cryogenic magnetometer or a spinner 1999; JNOC, 1997) are summarized in Table 1. As for magnetometer（ (Natsuhara-Giken SMM-85) ）settled the MITI Umaoi borehole, sampling depths are in a magnetically shielded laboratory at Kyoto converted to two-way travel time based on vertical University. seismic profiling (VSP), and shown on the interpreted profile in Fig. 3. 4.2.Demagnetization test 4.1. Sample preparation We conducted progressive thermal demagnetiza- tion (PThD) test in order to isolate stable compo- All samples of the Eniwa SK-1 and the Oligocene nents of remanent magnetization. PThD test was samples (UM2-1~6) of the MITI Umaoi were performed, l ）to 690 °C in air, using a non- dn successfully oriented based on correlation between inductively wound electric furnace with an internal directional markers (bedding plane and fracture) on residual magnetic field less than 10 nT. Demagnet- core surfaces and side-wall images by borehole ization was interrupted when specimens were broken logging. Fracture patterns upon the Eniwa SK-1 cores into numerous pieces because of repeated heating were identical with those appeared on the logging and cooling procedures. Table 1 Description of core samples obtained from the southern central Hokkaido Well name Core Depth (m) ID (core-horizon) Lithology (Formation) N Age (method) Eniwa SK-1 1 4577.71-0.78 EN1-1 mudstone 3 143-141 Ma (radiolaria) 4577.89-0.96 EN1-2 mudstone 7 143-141 Ma (radiolaria) MITI Umaoi 2 2377.28-0.48 UM2-1 siltstone (Minami-naganuma) 12 L. Oligocene (dinoflagellate) 2 2378.60-0.80 UM2-2 siltstone (Minami-naganuma) 11 L. Oligocene (dinoflagellate) 2 2379.20-0.40 UM2-3 siltstone (Minami-naganuma) 12 L. Oligocene (dinoflagellate) 2 2380.30-0.50 UM2-4 siltstone (Minami-naganuma) 9 L. Oligocene (dinoflagellate) 2 2381.68-0.88 UM2-5 siltstone (Minami-naganuma) 9 L. Oligocene (dinoflagellate) 2 2383.08-0.26 UM2-6 siltstone (Minami-naganuma) 12 L. Oligocene (dinoflagellate) UM6-1 6 4851.56-0.63 gabbro 3 102, 101 Ma (K-Ar) 7 5018.27-0.37 UM7-1 gabbro 1 128 Ma (K-Ar); 95 Ma (FT) 7 5018.41-0.51 UM7-2 gabbro 5 128 Ma (K-Ar); 95 Ma (FT) N is number of specimens for paleomagnetic measurements. Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 89 Fig. 5 shows typical results of PThD for the MITI 1988） was adopted to determine directions of Umaoi Oligocene sedimentary rocks. Some horizons characteristic remanent magnetization (ChRM). (a) are characterized by overlapping TuB spectra of Other horizons (b) had univectorial demagnetization primary and secondary components, then remagne- trend apart from recent geomagnetic field direction, tization circle fitting (McFadden and McElhinny, then principal component analysis (Kirschvink, 1980) was adopted for ChRM calculation. As a result, we obtained four normal and reversed mean a) MITl Umaoi Core 2, Horizon 3, PThD (in-situ) directions for the Late Oligocene Minami-naganuma W_Up N Formation (Fig. 5c, Table 2). All of them are characterized by counterclockwise deflection in S HN 3140 400 declinations after tilt-correction. 370 400 Fig. 6 presents typical results of PThD for the 220 Cretaceous rocks of the Eniwa SK-1 and MITI 190Q Umaoi. Reddish mudstones from the Eniwa SK-1 160 Q130 (EN-series) preserve three distinct components of J/Jo NRM. A high-temperature component (600<unblock- Q100 ing temperatures (TuB)<690 °C） with easterly reversed directions was isolated, as was an intermedi- EDown 25℃℃ 400℃ Jo=5.89E-8 Am2 ate-temperature component (400<TuB<580 °C) with b) MITI Umaoi Core 2, Horizon 4, PThD (in-situ) westerly normal directions. A low-temperature com- WUp HN ponent (150<TuB<400 °C) was identified, and its SF N 400 northerly direction similar to that of the present 340 370 geomagnetic field suggests that the component is originated from thermoviscous remanent magnetiza- 220 tion (TVRM) acquired during the latest Brunhes Q190 Q160 normal polarity chron. 130 Two distinct components were isolated from NRM of the MITI Umaoi gabbro samples (Fig. 0100 6b; UM-series). A high-temperature component (520<TuB<580 °C) with shallow inclinations was EDownO25°℃ Jo=6.12E-8 Am 2 isolated, whereas a low-temperature component 400℃ c) MITI Umaoi C2 Minaminaganuma F. (untilted) (150<TuB<500 °C) with normal inclinations was N identified. As mentioned before, cores of this T borehole were not initially oriented, thus origin of Fig. 5. (a, b) Results of progressive thermal demagnetization (PThD) for the Late Oligocene core samples in the MITI Umaoi borehole. Directional data are shown up to 400 °C because these samples became quite unstable and acquired secondary remanent magnetization after higher levels of PThD. Vector-demagnetization diagrams are drawn on the left. Unit of the coordinates is bulk remanent intensity. Solid and open circles are projections of vector end-points on the horizontal and N-S vertical planes, respectively. Equal-area projections and normalized intensity decay curves are on the right. (O) Plotted on the lower hemisphere of the projections. Dm: -45.8° Numbers attached on the symbols depict levels of PThD in degrees Im: 40.9° Q95:25.2° Celsius. (c) Untilted characteristic mean directions from four K: 14.3 horizons have westerly deflection in declinations. Solid and open circles are plotted on the lower and upper hemispheres of the projection, respectively. Dotted ovals are 95% confidence limits. 90 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 Table 2 Paleomagnetic directions of the Minami-naganuma Formation in MITI Umaoi, southern central Hokkaido Horizon Facies Method DMG (°C) D (°) I() Dc () Ic () 0.95 (°) K N 中 UM2-2 siltstone RC 100-340 158.2 50.1 170.3 -42.4 2.1 960.2 8 70.0 11.7 siltstone 74.7 -35.1 3.4 34.0 UM2-3 RC 100-340 138.2 118.8 367.5 8 47.1 UM2-4 siltstone PCA 130-400 34.7 70.9 -40.4 50.1 4.9 128.2 8 55.9 44.7 UM2-5 siltstone PCA 130-420 -53.5 71.1 -64.0 27.7 2.9 373.0 8 28.9 45.0 RC and PCA are the remagnetization circle and principal component analysis, respectively, for determination of the mean directions. DMG is demagnetization range for calculation by RC/PCA method; D and I are in situ site-mean declination and inclination, respectively; Dc and Ic are untilted site-mean declination and inclination, respectively; α9s is radius of 95% confidence circle; K is the Fisherian precision ns s ( s u n directions, respectively. the NRM components will be discussed in Section from MITI Umaoi (UM6), soft coercivity fraction 4.4 based upon magnetic mineralogy described in the (<0.12 T) with TuB spectrum up to 580 °C is next section. dominant. Hence, the stable primary NRM is to be solely carried by titanomagnetite. 4.3.Magnetic mineralogy of the Cretaceous samples 4.4.Origin and correction of NRM components of the In order to identify ferrimagnetic minerals of the Cretaceous samples Cretaceous samples characterized by multicomponent remanence, stepwise acquisition experiment of iso- We determined the direction of each magnetic thermal remanent magnetization (IRM) was per- component using a three-dimensional least squares formed on AF-demagnetized specimens in direct analysis technique (Kirschvink, 1980). Fig. 8 presents magnetic fields up to 2 T (Fig. 7a). As for the mean directions of isolated NRM components. As for specimen from Eniwa SK-1 (EN), IRM intensity is the Eniwa SK-1 (EN), high- and intermediate-temper- characterized by initial rapid increase and gradual ature components are nearly antipodal, showing acquisition up to 2 T, indicating presence of magnetite counterclockwise deflections in declinations. Because and a high-coercivity mineral. On the other hand, the sample consists of reddish siliceous mudstone IRM intensity of the specimen from MITI Umaoi deposited in deep marine environment, high-temper- (UM) almost saturates in applied fields lower than 0.2 ature component carried by hematite may be post- T, and magnetite is the dominant contributor of stable depositional and prefolding magnetization as remanent magnetization. previously reported for cherts on the Eurasian margin Next, we executed PThD of composite IRMs for (Oda and Suzuki, 2000; Shibuya and Sasajima, 1986). the same specimens. Based on the procedure proposed Thus, we adopt untilted direction of the high-temper- by Lowrie (1990), composite IRMs were imparted by ature component for the tectonic discussion. applying direct magnetic fields of 2, 0.4 and then 0.12 Because low-temperature component of the MITI T onto the specimens in three orthogonal directions. Umaoi (UM) is characterized by inclination consistent As for the specimen from Eniwa SK-1 (EN1), decay with that of Earth's dipole field (62°) at the site curve of the IRM components through PThD test (see latitude, we assume that the component is recent Fig. 7b) indicates that the dominant magnetic phase is TVRM in origin and has in-situ northerly declination. the high (2~0.4 T) coercivity fraction with broad This assumption is supported by the IRM experi- spectrum of TuB up to 680 °C. Hence, carrier of the ments, showing that magnetite is the sole contributor high-temperature component of NRM is hematite. for stable NRM components. As depicted in Fig. 9, Both medium (0.4~0.12 T) and soft (<0.12 T) theoretical contour for thermoviscous remagnetization coercivity fractions of less amount are identified, for magnetite (Walton, 1980; Middleton and Schmidt, and we interpret that the fractions are mainly carried 1982) indicates that natural heating at ~150 °C, which by larger (MD-size) grains of hematite because they is estimated by_ temperature logging (BHT), for have TuB spectra up to 680 °C. As for the specimen duration of ~10° years, approximately corresponds 30° 670 548 680 a) EN1-2-3 PThD (in-situ) b) UM7-1-1 PThD (arbitrary north) Y. Itoh et al. / Tectonophysics 395 W_Up W_Up 2 400 20°℃ 50-580 350Q 300 100 20°℃ HN 250 150 600 200 + 150 0 bown 540-680 400 520 560 100 400 5(2005) 81-97 450 20°℃ 580 20℃8300 N "Down Jo=3.07E-7Am2 EDown 150 500℃ Jo=2.32E-5Am2 Fig. 6. Results of progressive thermal demagnetization (PThD) for the Cretaceous core : samples in the Eniwa SK-1 (a) and MITI Umaoi (b) boreholes. Vector-demagnetization diagrams are drawn on the left. Unit of the coordinates is bulk remanent intensity. Solid and open circles are projections of vector end-points on the horizontal and vertical planes, respectively. Vertical planes are in N-S and arbitrary directions for the Eniwa SK-1 and MITI Umaoi, respectively. Equal-area projections and normalized intensity decay curves are on the right. Solid and open circles are plotted on the lower and upper hemispheres of the projections, respectively. Numbers attached on the symbols depict levels of PThD in degrees Celsius. 6 92 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 a) IRM acquisition oriented directions of this component. As the sample 1.0 was taken from a basement block with gentle dips on the seismic record (Fig. 3), the in situ direction is 80 adopted for the tectonic discussion. 0.6 5. Discussion —-EN1-2-6 0.2 --UM6-1-2 Mean directions of the ChRM components of the 0.0 studied two boreholes are listed in Table 3 together 0 500 1000 1500 2000 with surface paleomagnetic data obtained from south- magnetic field (mT) ern central Hokkaido (Kodama et al., 1993). Although b) PThD of orthogonal IRMs suffered drilling-induced remagnetization (Kikawa, EN1-2-6 —Hard (2.0T) 1993), our data do not show high inclination values -Medium (0.4T) (~80°) suggestive of such secondary origin. There- 15 - Soft (0.12T) E fore, we regard that the paleomagnetic data are immune from drilling disturbance in the following lent discussion. In order to understand evolution of the deformation front, we argue tectonic implication of the paleomagnetic information and an organic matur- IRM 5 ity modeling (MM) of borehole samples in the following sections. 100 200 300 400 500 600 700 5.1.Extent of the Cretaceous volcanic arc and its temperature (°℃) rotationalmotion UM6-1-2 - Hard (2.0T) *- Medium (0.4T) 500 Basement rocks of the MITI Umaoi consist of 0-Soft (0.12T) E gabbro and metabasalt, which are correlated with igneous Kumaneshiri Group in the Rebun-Kabato belt (Fig. 1; Kato et al., 1990). As the igneous rocks 300 show geochemical affinity with arc-volcanism 200 (Nagata et al., 1986; Ikeda and Komatsu, 1986), distribution of the unit corresponds to extent of the early Cretaceous arc associated with hypothetical westward subduction (Niida and Kito, 1986). 0 100 200 300 400 500 600 700 Among exploration boreholes, MITI Sorachi and temperature (°℃) MITI Nanporo (Fig. 2) confirmed igneous base- Fig. 7. (a) Progressive acquisition of isothermal remanent magnet- ments similar to the Kumaneshiri Group (Kato et ization (IRM) for core samples in the Eniwa SK-1 (EN) and MITI al., 1990). Thus, the Ishikari lowland seems to be Umaoi (UM) boreholes. (b) Thermal demagnetization curves of underlain by the Cretaceous igneous rocks, which orthogonal IRMs for samples in the Eniwa SK-1 (upper) and MITI Umaoi (lower) boreholes. extend to the Oshima belt of western Hokkaido (Fig. 1; Kato et al., 1990). to maximum TuB ranges of the low-temperature From the Oshima belt, Otofuji et al． (1994) component in laboratory. On such ground, the high- reported Paleogene NRM directions with large temperature component is converted into in-situ westerly deflections, and interpreted that the belt coordinates (Fig. 8e, e' ). Shallow reversed inclination belonged to northeast Japan rotated counterclock- and counterclockwise deflection characterize the wise according to the back-arc spreading in the a) Low temperature Dm: -20.4 b) Middle temperature Dm: -80.1 c) High temperature N N Dm:131.3 component (in-situ) N Im: 67.4 component (in-situ) Im: 61.2 component (in-situ) Im: -48.0 095: 9.5 2'9 :960 095: 7.2 2 N:4 b') Middle temperature c') High temperature EN1-1 Dm: -56.9 Dm: 147.5 K:95.0 component(untilted)N component (untilted) N upperhem Im: 71.0 Im: -51.8 R:3.9684090 lowerhem 095: 6.7 095:7.2 K Itoh et al. /Tectonophysics 395 N: 7 N: 8 5(2005) 81-97 K: 83.1 upperhem K: 59.6 R:6.9277906 Oupper hem lowerhem R:7.8825136 lower hem d)Lowtemperaturecomponent e)Hightemperaturecomponent e')Hightemperaturecomponent (arbitrary north) N Dm:-126.0 (arbitrary north) N Dm:-48.8 (in-situ) N Dm: 77.2 Im: 62.3 Im: -26.6 Im: -26.6 α95:12.0 095: 7.2 095: 7.2 2 + M7- N:6 N:6 N:6 K: 32.2 K: 86.9 K: 86.9 Oupperhem upperhem R:5.8449101 ●lower hem R:5.9424815 lower hem R:5.9424815 ●lower hem Fig. 8. Mean magnetic directions obtained from the Eniwa SK-1 (EN) and MITI Umaoi (UM) boreholes. Solid and open circles are on the lower and upper hemispheres of the equal- area projections, respectively. Dotted ovals are 95% confidence limits. 6 94 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 1E+9 1E+7 in-situtemperaturerangeduringthe 1E+5 latestBrunhesChronestimatedfrom BHTlogging data in MITI "Umaoi" 1E+3 1E+1 1E-1 1E-3 1E-5 1E-7 100 200 300 400 500 600 Temperature (o℃) Fig. 9. Thermoviscous remagnetization contours (thin solid curves) after Middleton and Schmidt (1982) based on the SD relaxation theory (Walton, 1980). Shaded oval depicts plausible remagnetization temperatures reached in nature, estimated from the logging data of botom-hole temperatures conducted on the MITI Umaoi (JNOC, 1997). Horizontal arrow shows the range of laboratory unblocking temperatures (TuB) of "low-temperature component? residing in the UM core sample (gabbro). Neogene. As shown in Table 3, the Cretaceous Lallemand and Jolivet, 1986; Kimura, 1996), which is paleomagnetic data in our study area are charac- postulated based on structural geology and tectonic terized by westerly deflections, although quantitative framework. Late Oligocene to early Miocene right- evaluation is difficult for the small number of data lateral shear is detected on deformation of sedimen- with ambiguities in correction of tectonic tilting as tary rocks (Kusunoki and Kimura, 1998). Oblique stated in Section 4.4. Thus, the downthrown side of collision between the North American (Okhotsk) and the floor thrust of the Hidaka collision front may be an Eurasian Plates, and spreading of the Kuril Basin extension of the Cretaceous igneous belt of northeast (Kusunoki and Kimura, 1998) are assumed cause of Japan. the transcurrent motion. Based on en echelon arrange- ment of anticlines of oil fields around MITI Karumai 5.2.Rotational deformation of the collisionfront borehole (Fig. 2), Oka (1986) suggested that the wrench deformation had been lingering on the western Surface paleomagnetic data obtained from southern flank of the Hidaka Collision Zone until the middle central Hokkaido (Table 3; Kodama et al.,1993) is Miocene. Thus, long-standing right-lateral shearing is indicative of clockwise rotational motion. Considering widely accepted as structural trend in southern central ages of the measured rock units, clockwise rotation is Hokkaido. Takeuchi et al. (1999) proposed a model of assigned to the early to middle Miocene, and ball-bearing clockwise motions of whole lithosphere interpreted as a result of right-lateral shear (e.g., blocks (Fig. 1b). Table 3 Paleomagnetic data reported from the southern central Hokkaido Area/Well Geologic age Remarks Declination (°) Inclination (°) 0.95 (°) N Yubaria Oligocene-Miocene E(U) -155.0 -64.0 14.0 8 MITI Umaoi Late Oligocene E(U) -45.8 40.9 25.2 147.5 Eniwa SK-1 Cretaceous W(D) -51.8 1 MITI Umaoi Cretaceous W(D) 77.2 26.6 1 N is number of sites. a Data from Kodama et al. (1993). W(D) and E(U) are western (downthrown) and eastern (upthrown) sides of the Umaoi Fault, respectively. Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 95 However, our Oligocene paleomagnetic data convergence history in larger scale and resultant obtained from the MITI Umaoi show counterclock- mountain-building process in Hokkaido is not fully wise sense of rotations, opposite from the almost understood, the recent tectonic episode may enhance contemporaneous surface directions (Kodama et al., contraction and thrust activity on the deformation 1993). In spite of fairly large directional scatter front. (Table 3), comparison of the two data-sets indicates In an effort to determine the duration of thrusting, significant relative rotation of 70.8° with △R (Beck, we conducted a one-dimensional maturity modeling 1980） of 47.9°.It is noted that both of these at the MITI Umaoi borehole for the downthrown datasets belong to the eastern (upthrown) side of the rock unit of the floor thrust in Fig. 3 with the LLNL floor thrust of the deformation front. Because single kinetic model (Sweeney and Burnham, 1990). Based fault large enough to compensate the differential on diatom assemblages of the growth strata (Kurita rotation has not been reported between these areas and Yokoi, 2000), oldest initiation of thrusting is (see Fig. 2), the present result suggests more assigned to ca. 11 Ma. Youngest initiation of the complicated rotational process in the transpressional thrust is set at 2 Ma, since TuB range of the TVRM deformation zone. Paleomagnetic interpretation component (Fig. 9) residing in the basement rock linked with detail fault assessment based on implies recent burial. For the both scenarios, structural geology is required, as attempted by Mino maturity levels of organic matters are calculated et al. (2001) for northeast Japan arc. based on burial histories of the downthrown block. Fig. 10 presents the results of modeling for the 5.3.Recent E-W contraction on the collision front Eocene Poronai Formation together with measured maturity levels of vitrinite (%Ro). In the formation, Since the late Neogene, southern central Hok- maturitylevels shown by %Ro rise as burial depth kaido has been a site of arc-arc collision as a result increases, which implies contamination from differ- of westward migration of the Kuril fore-arc sliver ent horizons is not serious. Obviously, the youngest (Fig. 1; Kimura and Tamaki, 1985; Kimura, 1996). scenario accords with the maturation trends in Quaternary E-W contraction is detected on active nature. Therefore, we conclude that the E-W fault distribution (Ikeda et al., 2002) and geodetic contraction along the collision front has been data (Tada and Kimura, 1987). Although the plate activated in the Quaternary. The continental collision 3600 3800 + measured Ro in Poronai F. (w) 4000 calculated maturity in case thrusting since 2 Ma calculated maturity in case 4200 thrusting since 11 Ma 4400 0.1 Maturity (%Ro) Fig. 10. Organic maturity levels of the Poronai Formation (Eocene) in the MITI Umaoi borehole. Vitrinite reflectance data (%Ro) measured on cuttings are shown by dots. Maturity levels calculated on the basis of a kinetic model are drawn as thin lines, which represent possible scenarios of thrusting of the Umaoi Fault and burial history of the downthrown-side block containing the geochemical sampling interval (MM in Fig. 3). 96 Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 in central Hokkaido is an ongoing tectonic event and an anonymous referee helped to improve this with westward propagation. manuscript. 6. Conclusions References (1) Downthrown block of the floor thrust of the arc- Arita, K., Toyoshima, T., Owada, M., Miyashita, S., Jolivet, L., 1986. arc collision front in southern central Hokkaido Tectonic movements of the Hidaka metamorphic belt, Hokkaido, is underlain by the Cretaceous igneous rocks Japan. Monogr. Assoc. Geol. Collab. Jpn. 31, 247-263. extending to northeast Japan. Westerly deflec- Beck Jr., M.E., 1980. Paleomagnetic record of plate-margin tectonic processes along the western edge of North America. J. Geophys. tions in ChRM directions of the subsurface Res. 85, 7115-7131. Cretaceous samples may be related to counter- Geological Survey of Japan, 2002. Geological Maps of Japan clockwise rotation of northeast Japan during the 1:200,000 (Images) Ver. 2.0. Geological Survey of Japan, Neogene back-arc opening. Tsukuba. Rotational motion of the upthrown block is Ikeda, I., Komatsu, M., 1986. Early Cretaceous volcanic rocks of controversial. Late Oligocene Minami-naga- Rebun Island, north Hokkaido, Japan. Monogr. Assoc. Geol. Collab. Jpn. 31, 51-62. numa Formation in the MITI Umaoi borehole Ikeda, Y., Imaizumi, T., Sato, H., Togo, M., Hirakawa, K., consistently exhibits westerly deflections in Miyauchi, T. (Eds.), 2002. Atlas of Quaternary Thrust Faults untilted ChRM directions, whereas surface in Japan. University of Tokyo Press, Tokyo. 254 pp. Ishiwada, Y., Aida, H., Atake, M., Araki, N., Iijima, A., Ikeda, A., coeval samples have easterly deflected ChRM. Okuda, Y., Kikuchi, Y., Kojima, K., Saito, T., Sato, Y., Tanaka, A previous model of ball-bearing clockwise S., Tono, S., Hirayama, J., Honza, E., Miyazaki, H., Morishima, motions of crustal blocks cannot account for H., Harada, Y., 1992. Oil and Natural Gas Resources in Japan. the reality. As a single fault, compensating such Natural Gas Exploration Association, Tokyo. 520 pp. differential motions has not been reported, closer Ito, T., 2000. Crustal structure of the Hidaka collision zone and its investigation on fault architecture and rotational foreland fold-and-thrust belt, Hokkaido, Japan. J. Jpn. Assoc. Pet. Technol. 65, 103-109. motion is required to understand deformation JNOC (Japan National Oil Corporation), 1996. Report for the mode of the arc-arc collision zone. Offshore Geophysical Survey Iburioki-Senkaiki, 1995 Fiscal (③) Duration of the floor thrust activities of the Year Japan National Oil Corporation, Tokyo, 35 pp. collision front is estimated utilizing the maturity JNOC (Japan National Oil Corporation), 1997. Report for the levels of organic matters in the downthrown Geological Study of MITI Umaoi Borehole, 1996 Fiscal Year. Japan National Oil Corporation, Tokyo, 61 pp. block. Vitrinite reflectance trend of the Eocene Kato, M., Katsui, Y., Kitagawa, Y., Matsui, M. (Eds.), 1990. Poronai Formation within the MITI Umaoi is Regional Geology of Japan: Part 1. Hokkaido. Kyoritsu best matched with a burial history assuming Shuppan, Tokyo. 337 pp. thrust activities since 2 Ma. Therefore, the E-W Katsumata, K., Wada, N., Kasahara, M., Okayama, M., Ichiyanagi, contraction along the collision front has been M., Ishikawa, H., Takada, M., Cho, I., Umino, N., Okada, T., Nakamura, A., Hori, S., Tachibana, K., Kono, T., Nida, K., activated in the Quaternary. Hashimoto, K., Ito, Y., Igarashi, T., Nakajima, J., Asano, Y., Ito, A., Uchida, N., Soda, Y., Ujikawa, H., Hasemi, A., Demachi, T., Hirata, N., Urabe, T., Sakai, S., Ide, S., Ogino, I., Seto, N., Acknowledgements Sakai, K., Hashimoto, S., Haneda, T., Yamanaka, Y., Miura, K., Hagiwara, H., Kobayashi, M., Inoue, Y., Tagami, K., Nakagawa, S., Tsuda, K., Matsubara, M., Tada, T., Aoyama, H., Matsu- The authors thank METI (Ministry of Economy, zawa, T., Zhao, Y., Yamazaki, F., Yamada, M., Sasaki, Y., Trade and Industry, Japan), JNOC (Japan National Hiramatsu, Y., Saiga, A., Komori, T., Umeda, Y., Ito, K., Oil Corporation) and JAPEX for the permission Koizumi, M., Wada, H., Hirano, N., Nishida, R., Matsushima, to publish this work. We acknowledge the follow- T., Uehira, K., Ooshima, M., Hirano, S., 2002. Distribution of ing persons: S. Yokoi for assistance during core hypocenters and focal mechanisms in and around the Hidaka arc-arc collision zone revealed by a dense temporary seismic sampling; H. Kurita for thoughtful discussion; N. network. Bull. Earthq. Res. Inst. 77, 199-223. Ishikawa for the use of paleomagnetic laboratory Kazuka, T., Kikuchi, S., Ito, T., 2002. Structure of the foreland fold- at Kyoto University. Reviews by R. Sutherland and-thrust belt, Hidaka Collision Zone, Hokkaido, Japan: re- Y. Itoh et al. / Tectonophysics 395 (2005) 81-97 97 processing and re-interpretation of the JNOC seismic reflection Middleton, M.F., Schmidt, P.W., 1982. Paleothermometry of the profiles “Hidaka’ (H91-2 and H91-3). Bull. Earthq. Res. Inst. Sydney Basin. J. Geophys. Res. 87, 5351-5359. 77,97-109. Mino, K., Yamaji, A., Ishikawa, N., 2001. The block rotation in the Kikawa, E.，1993. Drilling-induced remanent magnetization Uetsu area, northern part of Nigata Prefecture, Japan. Earth in gabbroic drill cores. J. Jpn. Soc. Marine Surv. Technol. 5, Planets Space 53, 805-815. 11-19. 4.53 Kimura, G., 1996. Collision orogeny at arc-arc junctions in the the Kabato Mountains: the age and nature as an Early Japanese Islands. The Island Arc 5, 262-275. Cretaceous volcanic arc. Monogr. Assoc. Geol. Collab. Jpn. 31, Kimura, G., Tamaki, K., 1985. Tectonic framework of the Kuril arc 63-79. since its initiation. In: Nasu, N., Uyeda, S., Kushiro, I., Niida, K., Kito, N., 1986. Cretaceous arc-trench systems in Kobayashi, K., Kagami, H. (Eds.), Formation of Active Ocean Hokkaido. Monogr. Assoc. Geol. Collab. Jpn. 31, 379-402. Margins. Terra Publishing, Tokyo, pp. 641-676. 2.26 Kirschvink, J.L., 1980. The least-squares line and plane and the red bedded chert of the Inuyama area, central Japan. J. Geophys. analysis of palaeomagnetic data. Geophys. J. R. Astron. Soc. 62, Res. 105, 25743-25767. 699-718. Oka, T., 1986. Distribution and tectonic evolution of Late Cenozoic Kodama, K., Takeuchi, T., Ozawa, T., 1993. Clockwise tectonic basins in Hokkaido. Monogr. Assoc. Geol. Collab. Jpn. 31, rotation of Tertiary sedimentary basins in central Hokkaido, 295-320. northern Japan. Geology 21, 431-434. Otofuji, Y., Kambara, A., Matsuda, T., Nohda, S., 1994. Counter- Koike, K., Machida, H. (Eds.), 2001. Atlas of Quaternary Marine clockwise rotation of Northeast Japan: paleomagnetic evidence Terraces in the Japanese Islands. University of Tokyo Press, for regional extent and timing of rotation. Earth Planet. Sci. Lett. Tokyo. 105 pp. with three CD-ROMs and two oversize maps. 121,503-518. Komatsu, M., Miyashita, S., Maeda, J., Osanai, Y., Toyoshima, T., 5.18 1983. Disclosing of a deepest section of continental-type crust case study in the Inuyama area, central Japan. J. Geophys. Res. up-thrust as the final event of collision of arcs in Hokkaido, 91,14105-14116. North Japan. In: Hashimoto, M., Uyeda, S. (Eds.), Accretion Suzuki, N., Kurita, H., Takashima, R., 1999. Earliest Creta- Tectonics in the Circum-Pacific Regions. Terra Publishing, ceous radiolarian assemblages from a deep borehole section Tokyo, pp. 149-165. in the southern Ishikari Plain, central Hokkaido, and their Komatsu, M., Osanai, Y., Toyoshima, T, Miyashita, S., 1989. implications. Abstract of 106th Annual Meeting of the Geo- Evolution of the Hidaka metamorphic belt, Northern Japan. In: logical Society of Japan. Geological Society of Japan, Tokyo, Daly, J.S., Cliff, R.A., Yardley, B.W.D. (Eds.), Evolution of Pp. 307. Metamorphic Belts. Geological Society, London, pp. 487-493. Sweeney, J.J., Burnham, A.K., 1990. Evaluation of a simple model Kurita, H., Yokoi, S., 2000. Cenozoic tectonic settings and a current of vitrinite reflectance based on chemical kinetics. Am. Assoc. exploration concept in southern central Hokkaido, northern Pet.Geol.Bull.74,1559-1570. Japan. J. Jpn. Assoc. Pet. Technol. 65, 58-70. Tada, T., Kimura, G., 1987. 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Itoh (2005) - Deformation mode in the frontal edge of an arc-arc collision zone.txt
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Thermal and geochemical structure of the
Uenotai geothermal system, Japan
Naoto Takeno*
Geological Survey of Japan, Tsukuba, Ibaraki, 305-8567, Japan
Received 12 January 1998; accepted 7 July 1999
Abstract
The Uenotai geothermal area is located in southern Akita prefecture of northern Honshu
Island. The Uenotai geothermal system is a liquid-dominated system with a central zone of
aquifer boiling. The two-phase reservoir has evolved from liquid in the natural state due to
exploitation. Gas composition of the vapor phase in the reservoir is nearly in equilibriumand correlates with the vapor fraction in the reservoir and with discharging steam quality.The marginal part of the Uenotai system has cooled with the drop in ground-water level.
The chemical characteristics of the geothermal water indicate mixing of the immature high
Cl source water with conductively heated or steam-heated shallow water or surface water,as well as boiling and steam gain. #2000 CNR. Published by Elsevier Science Ltd. All
rights reserved.
Keywords: Aquifer; Fluid chemistry; Fluid inclusion; Uenotai; Japan
1. Introduction
The Uenotai geothermal area is located in the southern part of Akita
Prefecture, northern Honshu, Japan, and consists of a liquid-dominated system
with a large boiling aquifer portion (Naka et al., 1987). Geothermal exploitation is
thought to have produced this e/C128ect on a liquid reservoir in natural state (Mimura
0375-6505/00/$20.00 #2000 CNR. Published by Elsevier Science Ltd. All rights reserved.
PII: S0375-6505(99)00062-0Geothermics 29 (2000) 257–277
* Tel.: +81-298-54-3708; fax: +81-298-54-3702.
E-mail address: g0442@gsj.go.jp (N. Takeno).
et al., 1995). Indeed, the Uenotai system di/C128ers from the vapor-dominated system
as defined by White et al. (1971). For example, large-scale vapor static conditionswere not confirmed during exploitation (Takenouchi, 1988; Naka and Okada,
1992). The evolution of the large boiling aquifer from the liquid reservoir may
provide insights on the vapor-dominated system. This is the aim of this reviewpaper on the Uenotai system.
There is an abundance of geothermal manifestations such as hot springs,
fumaroles and geothermal alterations in this area. A 27.5 MW geothermal powerplant has been operating since 1994, run by Akita Geothermal EnergyCorporation (AGECO), a subsidiary of Dowa Mining Company. Geothermalprospecting and related research have been conducted by AGECO, NEDO (New
Energy and Industrial Technology Development Organization), and researchers
from universities and the Geological Survey of Japan.
2. Geological setting
The Uenotai geothermal area is located 15 km west of the Quaternary volcanic
front, on the back-arc side near Kurikoma volcano. A gravity study (Suda et al.,1981) of this area showed that the Uenotai geothermal area is located within apolygonal gravity low measuring 18 27 km trending NW–SE along the basement
structure (Usuta et al., 1981, 1982) (Fig. 1). The overall gravity decreases steeply
on the margins of the basin, and flattens inside the basin. However, the basin is
subdivided into some local areas of low and high gravity anomalies. In the highgravity area, the pre-Tertiary basement crops out at the surface, whereas in thelow gravity area this basement rock was not encountered by drillholes of about2000 m depth. The isogal contours thus correspond closely to the horst andgraben structure of the pre-Tertiary basement (Usuta et al., 1981, 1982). Thehorsts commonly form topographic peaks such as that of Mt Takamatsu.
The Uenotai geothermal area is located on the shoulder of this local gravity
high, on the margin between a horst and graben. The horst is composed of pre-
Tertiary basement unconformably overlain by Miocene formations (Usuta et al.,1981, 1982). The pre-Tertiary basement is a complex of schist, ultra-basic rocks,gneiss and granite with a strong NW structural trend (Usuta et al., 1981, 1982).Part of the granite is mylonitized in a NW direction, called the Onikobe–Yuzawamylonite zone (Sasada, 1984, 1985). These basement structures are comparable tothe Japanese skeletal tectonic trend such as the Matsushima–Honjo structuralzone (Oide and Onuma, 1960).
The Miocene Doroyu Formation (Okada, 1976) consists of volcanic rocks and
marine sediments. Dacite tu/C128, welded tu/C128 of the Torageyama (locally called‘‘Minasegawa’’) Formation (Usuta et al., 1981, 1982), and lacustrine tu/C128aceous siltof the Sanzugawa Formation (Muto, 1965) overlie the Doroyu Formation inascending order. The thicknesses of the Torageyama and Sanzugawa Formationsare variable; they are thick in the graben and very thin or thin out over the horst.The Torageyama Formation overlies unconformably the Doroyu Formation (Fig.N. Takeno / Geothermics 29 (2000) 257–277 258
2). K–Ar ages of the Torageyama Formation show two volcanic episodes, the
younger ranging from 2.8 20.5 to 3.8 20.2 Ma, and the older from 5.7 21.4 to 6.0
20.8 Ma (Takeno, 1988). The lowermost part of the Sanzugawa Formation and
the Torageyama Formation, on the margin of the basin, include large talusbreccia that correspond to the steep decrease of isogal contours or the location of
faults (Usuta et al., 1981, 1982). Therefore, the Sanzugawa and Torageyama
Formations were deposited in a topographic depression surrounded by faults.
Fig. 1. Uenotai geothermal area and its geologic setting. Hot springs: 1, Oyu; 2, Oyasu; 3, Takimukai;
4, Tochiyu; 5, Doroyu; 6, Kawarage; 7, Arayu; 8, Yunomata; 9, Takanoyu; 10, Yunotai; 11,Akayumata. Geological map after Takeno (1988), hot spring and alteration zone after Kimbara (1988)and isogal contour in mgal, density: 2.4 g/cm
3, after Suda et al. (1981). A–B: trace of the geological
cross-section (Fig. 2).N. Takeno / Geothermics 29 (2000) 257–277 259
These characteristics are interpreted to indicate a volcano-tectonic structure such
as a caldera (Awata, 1984; Takeno, 1988; Ito et al., 1989).
The middle part of the Sanzugawa Formation in the caldera is also observed on
the horst, which forms the topographic peak of Mt Oyasu around which many
geothermal manifestations can be observed. The uplift of the horst thus occurredafter the deposition of the middle part of the Sanzugawa Formation, and hassome relationship with volcanism. Deposition of the Sanzugawa Formation wasfollowed by the latest volcanic activities associated with the eruption of theKabutoyama Formation, mainly welded tu/C128 at 0.32–0.34 Ma (fission-track dating;Nishimura et al., 1976), and the Takamatsudake volcanic unit at 0.20 Ma (fission-track dating; Nishimura et al., 1976). These units rest on Mt Oyasu and Mt
Takamatsu.
3. Hot springs and geothermal alteration
Surface alteration includes a bleached siliceous zone consisting of a-cristobalite,
tridymite, kaolinite, alunite and pyrophyllite. At Kawarage and Doroyu there areold sulfur mines. Both areas have siliceous zones that are surrounded by kaolinitealteration (Kimbara, 1988; Hayashi and Moriguchi, 1990).
The chemical composition of the hot springs in this area is shown in Fig. 3.
Many hot springs are Cl-rich, neutral pH springs, but the Arayu and Tochiyu hot
springs are acidic with SO
4derived from oxidation of H 2S in vapor. At Kawarage
there are HCl-type acidic (pH 1.35) hot waters. These acidic springs have thehighest Cl content (2011 mg/l) of the hot springs and well discharges in theUenotai area (Table 1). The sum of the molal concentration of Na, K and Ca isless than one fifth that of Cl. The rest of the Cl and SO
4is balanced by Al, Mg,
and Fe+2, which are minor components in neutral pH springs. The SO 4cannot
Fig. 2. Schematic cross-section of Uenotai geothermal area. Partly modified from Kuriyama (1985).N. Takeno / Geothermics 29 (2000) 257–277 260
account for all of these cations, indicating that HCl is needed for the reaction
with the rock to release Al, Mg, and Fe+2. Kawarage acidic alteration is thought
to have been caused by Cl-rich acidic hot springs (Abe et al., 1979), and the
contribution of a volcanic component including HCl to the Kawarage hot spring
was also noted by Nakamura et al. (1955) and Abe et al. (1979).
The3He/4He–4He/20Ne ratio of the fumarolic gas at Kawarage indicates that
most of the He is derived from the mantle (Kita et al., 1992). These observations
suggest that the Kawarage fumaroles and hot spring discharges are closely related
to volcanic activity.
Subsurface geothermal alteration was studied by Naka and Okada (1992). The
exploration wells revealed a clear zoning sequence of zeolite minerals, from
mordenite through laumontite to wairakite with increasing depth. Epidote occurs
in the deeper part, with wairakite. Fractures are filled with quartz, calcite and
zeolite. Illite, illite/smectite intra-stratified minerals, chlorite, and minor anhydriteare also present. Minerals characteristic of acidic alteration, such as kaolinite,
pyrophyllite and alunite, do not occur in the subsurface alteration of the Uenotai
area. The subsurface alteration minerals indicate a neutral-to-alkalic hydrothermal
environment, agreeing with the chemistry of the well discharges.
Fig. 3. Cl–HCO 3–SO 4diagram for the hot springs.N. Takeno / Geothermics 29 (2000) 257–277 261
Table 1
Chemical composition of hot springs (from Abe et al. (1979))
Oyu 1 Oyasu 2 Takimukai 3 Tochiyu 4 Doroyu 5 Kawarage 6 Arayu 7 Yunomata 8 Takanoyu 9
Temp ( 8C) 96.5 99.5 90.8 88.5 83.3 88.7 95.7 51.5 74.5
pH 8.9 8.7 8.4 3.7 5.5 1.4 2.8 7.1 7.6Cl (mg/l) 237 270 227 3.55 306 2012 8.51 1.77 654
SO
4 117 149 125 17.7 143 1411 162 0.41 49.8
HCOÿ
3 3.51 4.58 41.2 – 35.1 – – 458 65.6
CO2ÿ
3 51.0 37.5 9.00 – 0.00 – – 0.00 0.00
CO2 0.00 0.00 0.00 – 135 – – 73.0 6.54
H2S 2.22 2.22 1.33 0.09 7.53 0.04 3.10 tr. 0.09
Ca 6.88 17.5 16.3 1.63 60.8 130 6.19 26.3 26.3Na 227 245 190 5.03 110 56.3 9.50 138 400
K 18.8 23.5 20.5 0.45 35.0 42.5 2.75 4.75 52.5
Mg 0.00 1.69 0.16 1.10 38.1 117 2.88 0.86 0.94Fe
2+0.00 0.85 0.05 0.00 0.05 4.85 2.62 0.00 0.00
Al 0.00 2.22 0.00 0.00 0.02 314 10.7 0.00 0.00
NH+
4 1.47 0.81 0.34 0.56 25.4 5.70 1.42 0.27 0.20
SiO 2 207 206 156 9.00 107 222 96.0 91.9 104
HBO 2 10.7 10.0 8.34 0.00 0.95 46.5 0.24 1.19 8.81
TDS 927 1009 858 63.0 964 4568 362 563 1365
dDSMOW (-)ÿ64.5ÿ65.9ÿ62.6ÿ62.3ÿ58.0ÿ54.5ÿ52.8ÿ64.5ÿ64.2
d18OSMOW (-)ÿ10.05ÿ10.15ÿ9.85ÿ9.54ÿ8.04ÿ7.03ÿ6.82ÿ10.27ÿ10.33
T=H(TU) 0.98 20.17 – 14.1 21.7 – 39.2 25.2 – 22.3 22.7 – –N. Takeno / Geothermics 29 (2000) 257–277 262
4. Geothermal reservoir
The Uenotai geothermal area is located at the transition between the horst
(Okumiyasan–Oyasudake uplift zone) and graben (Kijiyama subsidence zone), onthe northern flank of the uplifted basement (Okada, 1976). The WNW–ESEDoroyu fault forms the boundary between these two zones on the southwestmargin of Uenotai.
Stratigraphic data of many exploration wells indicate the existence of stepwise
faulting of small blocks separated by WNW–ESE faults parallel to the Doroyufault and NE–SW cross-cutting faults (Naka and Okada, 1992).
The subsurface temperature distribution at 500 m below sea level (Fig. 4) shows
that the high temperature zone, above 300 8C, extends northwestward from the
southeast part of Uenotai, corresponding to the stepwise fall of the basement (Fig.5). This suggests that the heat source is located southeast of Uenotai, toward MtTakamatsu, which is the center of the regional geothermal manifestations.
The steam quality (i.e., percentage of steam in total discharge), is shown in Fig.
4. High quality wells (some steam-only or superheated steam), such as T-5, T-29,
T-41, T-42, T-44, and T-46, are located inside or near the 300 8C contour, and the
Fig. 4. Temperature distribution ( 8C) at 500 m below sea level. C zone includes high quality wells; M
zone includes low quality wells, the margin of the system. Compiled from Naka and Okada (1992).N. Takeno / Geothermics 29 (2000) 257–277 263
low quality wells are located around them. Many wells have ‘‘excess’’ enthalpy,
because total discharge enthalpy is greater than the enthalpy of saturated liquidwater at the temperature estimated from the quartz adiabatic geothermometer orNa–K–Ca geothermometer (Table 2).
Wells T-41 and T-42 produced superheated steam with temperatures greater
than 180 and 165 8C at the wellhead, respectively (Naka et al., 1987). Discharge
quality decreases outward from the area of steam-only wells, corresponding to thetemperature decrease.
Some wells are reported to have changed their quality from liquid–vapor
mixture to dry steam-only or superheated steam. For example, this occurred at T-21 after 4 months of discharge (Takeuchi, 1996), and at T-41, T-42 and T-46 after
2–3 months (Naka and Okada, 1992).
The logging temperature of the wells (Fig. 6) (Takenouchi, 1988) indicates that
Fig. 5. Elevation above sea level of the top of Doroyu Formation. Hatched area: outcrop of Doroyu
Formation. T-34, YO-7, etc. are wells. After Naka et al. (1987).N. Takeno / Geothermics 29 (2000) 257–277 264
Table 2
Properties of the Uenotai wellsa
KT-1 T-13 T-27 T-34 T-41 T-42 T-44 T-45 T-46 T-49 T-50 T-51
Year of starting to drill 1978 1974 1979 1978 1980 1980 1983 1985 1985 1986 1986 1986
Drilling length (m) 1310.0 1010.4 1314.0 1700.7 1191.0 1501.0 1405.0 1666.9 1710.6 1849.0 1665.0 1776.0
Max temp. ( 8C) 290 290 290 300 305 315 295 290 310 310 335 295
Wellhead press. (kg/cm2gauge) 0.5 1.3 0.02 0.4 6.0 6.0 6.0 6.0 6.0 6.0 6.0 6.0
Vapor (t/h) 2.8 14.2 0.6 1.9 28.2 54.6 27.6 42.0 31.9 22.2 5.1 60.7
Liquid (t/h) 2.5 21.1 1.5 3.3 0 0 0 45.0 0 4.9 2.7 27.8Discharge enthalpy (kJ/kg) 1641 1402 1065 1273 – 2093
b– 1692 2512c2387 2046 2112
Liquid enthalpy (kJ/kg) at Tqa 1130 976 1000 1071 – – – 1211 – 1285 939 1317
Steam fraction 0.528 0.402 0.286 0.365 1 1 1 0.483 1 0.819 0.654 0.686
Factor 1.44 1.30 1.04 1.13 – – – 1.44 – 3.88 1.79 2.12
aDischarge enthalpy is calculated from flow-rate and wellhead pressure, except for T-42 and T-46. Factor is calculated following the method of Seki
(1990) (see text).
bMenzies et al. (1990).
cPham et al. (1995).N. Takeno / Geothermics 29 (2000) 257–277 265
all wells are at hydrostatic pressure. Mimura et al. (1995) also suggested that the
pressure of the excess enthalpy wells of the Uenotai geothermal area is close to
hydrostatic. The geothermal reservoir of this district is therefore thought to have
been liquid-dominated before its exploitation. After exploitation the wells haveevolved to produce steam only in the high-temperature area.
The character of the wells was summarized by Naka et al. (1987). Steam is
produced from fractures, faults and lithological boundaries in the Doroyu
Formation and pre-Tertiary basement. The Pre-Tertiary basement, a complex ofgranite, schist and ultra-basic rocks, hosts many acid intrusives that occur in theDoroyu Formation. The productive feed points of the steam-only wells are mostlyassociated with fractures around acid intrusives in the pre-Tertiary basement.
Fig. 6. Homogenization temperature of fluid inclusion (triangles with bars) and logging temperatures of
the wells (solid lines). Vertical axis: depth (m), horizontal axis: temperature ( 8C). Boiling point curve
(broken line) is that of 1 wt % NaCl from the present water-level. Partly modified from Takenouchi
(1988).N. Takeno / Geothermics 29 (2000) 257–277 266
Numerical simulation of the natural state of the Uenotai reservoir before
production indicates a low permeability and the importance of fractured rock forfluid flow (Antunez et al., 1990). Fractures in the Uenotai system are also
emphasized from the study of surface geology (Tamanyu and Mizugaki, 1993) and
tectonic simulation (Mizugaki, 1991, 1992).
The Sanzugawa Formation occurs in the upper part of all wells except T-5, and
its average thickness is 400 m. The Sanzugawa Formation is largely tu/C128aceous silt,and its incompetency, and lack of lost circulation zones (Naka et al., 1987), resultsin the formation being an e/C128ective cap rock (Akibayashi et al., 1981; Tamanyuand Mizugaki, 1993).
This hydrological condition in and around the reservoir caused an insucient
recharge of cold water to be heated near the wells and to expand the boiling
aquifer around the steam-only wells.
PTS (pressure, temperature and spinner) logging during the 1988 production
test indicated multiple feed points of deep, steam-only inflow zones and shallowliquid inflow zones for T-41, T-42, T-44 and T-50 (Menzies et al., 1990). Thisimplies that the deep boiling reservoir of the Uenotai geothermal system isoverlain by a shallow liquid reservoir, and water in the shallow reservoir may beheated by steam rising from the deep boiling reservoir.
In the Kawarage subsidence area, the isotherms are depressed (Fig. 2) and local
temperature reversals are reported (Akibayashi et al., 1981; Robertson-Tait et al.,1990). A numerical reservoir simulation indicates that this feature is caused bydownflow of recharge water (Akibayashi et al., 1981). In this regard, theKawarage and Kijiyama subsidence areas are important recharge areas of thisgeothermal system.
5. Chemistry of the geothermal water
The chemical composition of the geothermal water has been compiled from
published data (Table 3). A Na–K–Mg diagram (Fig. 7) shows that many well
discharges plot along a mixing line between KT-1 and T-49. T-13 and T-27 ploton the low-temperature side of this mixing line, in agreement with their locationon the margin of the low subsurface temperature area (Fig. 4).
T-49 is the end member furthest from the full equilibrium line of all the data
plots along the line between KT-1 and T-49. Solutions far from the fullequilibrium line, such as T-49, will react with the rock in order to attainequilibrium, changing its composition toward KT-1. In this regard, the
composition of T-49 should be close to the parental water of the plots from T-49
to KT-1.
The exceptional characteristics of T-41, T-42L, T-44, and T-50 may be ascribed
to a low-temperature inflow from the shallow multi-feed points that have beendetected by PTS logging (Menzies et al., 1990).
The chemical composition of T-41 and T-44 plots in the field of immature
waters: these two waters are characterized by low Cl concentration (less thanN. Takeno / Geothermics 29 (2000) 257–277 267
Table 3
Chemical composition of fluid discharging from the Uenotai wellsa
KT-1bT-13bT-27bT-34bT-41cT-42bT-42LcT-44cT-45cT-46cT-49cT-50cT-51c
Liquid phase
Sampling date 1983.1 1980.12 1980.7 1982.5 1988.10 1980.12 1988.11 1988.10 1988.11 1988.11 1988.12 1988.12 1988.11
pH of liquid phase 9.8 9.6 9.3 9.9 9.9 9.4 10.0 8.8 9.5 9.0 8.8 9.4 9.5Cl (mg/l) 464 665 464 240 151 367 343 15.6 278 779 885 307 326SO
4 108 126 101 21.0 130 76.0 59.0 135 2.5 232 46.0 54.0 7.8
H2S 12.2 8.8 20.6 13.2 8.17 24.7 78.7 6.43 1.1 4.79 1.27 65.9 3.31
HCO 3 360 46.0 94.0 140 60.0 36.0 321 56.0 0.0 41.0 118 1436 128
Ca 3.09 4.71 4.96 0.51 12.7 1.21 1.7 19.1 0.36 1.78 9.3 2.24 0.57Na 360 437 425 200 251 247 602 67 214 671 620 1048 240
K 64.2 54.4 50.1 43.4 38.3 58.8 11.6 23.5 42.9 126 124 21.3 46.8
Mg 0.04 0.02 0.08 0.04 3.94 0.13 0.08 0.79 0.03 0.47 1.6 0.35 0.11SiO
2 775 507 543 668 894 621 523 225 943 1561 1106 455 1174
Steam phase
Sampling date 1978.10 1977.10 1988.11 1984.10 1988.11 1988.11 1988.11 1988.11 1988.12 1988.12 1988.11Non-condensable
gas (vol.%)0.23 0.14 0.246 0.18 0.104 0.120 0.140 0.141 0.223 0.246 0.264
CO
2 84.9 85.1 73.8 88.2 73.7 70.2 80.0 73.4 84.1 87.9 86.2
H2S 10.3 7.90 9.75 5.90 12.4 13.7 13.3 10.7 6.55 2.14 6.69
H2 2.69 0.203 14.7 4.44 13.2 16.0 3.28 13.2 6.45 6.71 5.25
O2 0.005 0.056 n.d. tr. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
N2 2.02 6.36 1.24 0.43 0.31 0.00 3.13 2.68 2.81 3.01 1.42
CH 4 0.0816 0.385 0.516 0.025 0.343 0.155 0.326 0.00 0.0859 0.260 0.453
Tqa(8C) 259 227 232 247 271 242 229 175 275 322 289 219 295
TNaKCa (8C) 252 227 222 275 217 276 140 238 275 279 261 149 270
Tkn(8C) 283 250 246 303 268 312 128 356 295 288 294 131 292
Tkm(8C) 218 226 191 200 113 187 136 122 205 192 167 131 181
aGas compositions of T-27 and T-34 are not available. Liquid samples of T-41, T-42, T-44, and T-46 were collected before changing to vapor-only pro-
duction.
bNaka et al. (1987).
cNaka and Okada (1992).N. Takeno / Geothermics 29 (2000) 257–277 268
about 150 mg/l, the most dilute in Table 3), and relatively high SO 4concentration
(the highest except for T-46). Furthermore, T-41 and T-44 have become steam-only wells. These characteristics suggest that the initial discharge water of T-41
and T-44 was steam condensate including SO
4oxidized from H 2S, and was not
fully equilibrated with the rock. The water originated from the shallow part of thewell, and from the upper multi-feed points of the well as indicated by the PTSlogging.
The chloride concentration–enthalpy diagram is shown in Fig. 8. The chloride
concentration of liquid from the excess enthalpy well was estimated by the method
proposed by Seki (1990). Seki (1990) introduced a correction factor for the
estimations of the chemical composition of the reservoir fluid of excess enthalpywells. His equation requires a downhole temperature and a reservoir temperature
for steam enthalpy. However, in this study these temperatures are assumed to be
the same because they are not available. Moreover, the enthalpy di/C128erence ofsteam between 200 and 300 8C is small, such that this assumption does not lead to
significant error.
For estimates of reservoir temperature, we used the conventional quartz
Fig. 7. Na–K–Mg diagram of liquid from the Uenotai wells. Full equilibrium line and the region of
immature water are from Giggenbach (1988).N. Takeno / Geothermics 29 (2000) 257–277 269
adiabatic geothermometer and K–Mg geothermometer proposed by Giggenbach
(1988). Due to the rapid reaction of water with minerals, the K–Mggeothermometer is expected to indicate the present reservoir temperature.
The temperatures estimated by the two geothermometers are plotted for each
well. In this diagram, the steam-only wells were omitted. Assuming one parentwater, the other waters appear to be derived from it by dilution, either by coldsurface water or a conductively- and/or steam-heated shallow water. Note that T-49 plots near the parent water in this diagram.
The CO
2/H2S ratio and CO 2concentration in the total discharge for the wells
are plotted in Fig. 9. After the method used originally by Glover (1970) andHedenquist (1990), the liquid curve was modeled as 260 8C liquid of composition
Fig. 8. Enthalpy–chloride diagram of liquid discharging from wells.N. Takeno / Geothermics 29 (2000) 257–277 270
similar to T-49. A parent 290 8C liquid composition was calculated for single-step
separation, and coexisting vapor composition was also calculated for separation in208C steps. All wells except T-50 are enveloped by the two curves, and the gas
concentrations of the wells can be explained by a mixture of the liquid and vapor
phases.
The analytical concentration quotients K0
cfor gas equilibrium for
CH42H2OCO24H2
was proposed by Giggenbach (1980) as
K00
cxd,CO2x4
d,H2=xd,CH4
1ÿxgysBCO2, 2804
where xd,iis total discharge gas concentration, ysthe vapor fraction, xgthe gas
fraction, and BCO2,2 8 0 the CO 2distribution coecient for a 280 8C average
temperature. The log K0cvalues for Uenotai wells are plotted at the temperature
estimated by the quartz adiabatic and Na–K–Ca geothermometers (Fig. 10). Allwells except T-13 plotted in the field of vapor gain. High steam-quality wells T-41,T-42, T-44, T-49, and T-50 plot near the equilibration vapor-phase line. Wells
Fig. 9. CO 2(mmol/100 m)–CO 2/H2S ratio for total discharge compositions. Calculated liquid and
vapor curves are modeled assuming a 260 8C liquid of composition similar to T-46, with parent 290 8C
liquid composition calculated for single-step (10 8C interval) separation.N. Takeno / Geothermics 29 (2000) 257–277 271
with a higher fraction yagree with the higher steam quality of the well discharge
(Fig. 4).
The original vapor fraction ( y) in the reservoir was calculated using the
equation proposed by D’Amore and Celati (1983) and Henley (1984) (Fig. 11).The high ywells, T-41, T-42, T-50, and T-44, also plot near the vapor
equilibration line (Fig. 10). These wells are also high quality in the highsubsurface temperature zone (Fig. 4).
Based on the preceding data, two zones of wells have been identified (Fig. 4):
the C zone has the high quality wells located in the central part of Uenotai, above3008C at 500 m below sea level, while the M zone contains the low quality wells
located on the margins of the system, below 280 8C at 500 m below sea level. The
C zone includes wells T-41, T-42, T-44, T-46, and T-50, and the M zone includeswells KT-1, T-5, T-13, T-21, T-27, T-29, T-34, and T-51.
The C zone wells are characterized by a high steam ratio or steam only, in the
two-phase reservoir with high steam fraction. As discussed before, some of the Czone wells are a/C128ected by a shallow, low-temperature fluid that causes low fluidmaturity or low temperature geothermometry. This shallow fluid is the condensateof steam rising from the deeper two-phase reservoir. As discussed, the gas
Fig. 10. Plot of temperature [quartz (open circle) and Na–K–Ca (dot) geothermometer temperatures] vs
log of K0c.N. Takeno / Geothermics 29 (2000) 257–277 272
Fig. 11. Original vapor fraction yin the reservoir, calculated after D’Amore and Celati (1983) and Henley (1984).N. Takeno / Geothermics 29 (2000) 257–277 273
composition of CH 4,H 2O, CO 2, and H 2in the vapor phase in the reservoir
indicates near-equilibrium conditions. T-49 liquid is close to the composition ofthe source fluid. The other fluids are derived from this source fluid mainly by
dilution, boiling and steam gain.
6. Fluid inclusion study and thermal history
Takenouchi (1988) discussed the relationship between observed temperature
and homogenization temperature ( T
h) of the fluid inclusions (Fig. 6) in calcite,
quartz, and wairakite. Many inclusions are liquid-dominant, and lack boilingcharacteristics (i.e., there are no vapor-rich inclusions). However, primaryinclusions in a quartz vein at 1251.7 m depth in well T-21 contain halite, sylvite,liquid, and vapor of variable ratios (Takenouchi, 1988). This indicates that therewere local boiling conditions and a high NaCl concentration in the Uenotaisystem. Local boiling toward dryness for the brine-filled fluid inclusions such as atUenotai are also suggested in the Valles caldera (Sasada and Go/C128, 1995), and inthe Broadlands-Ohaaki geothermal system (Simmons and Browne, 1997).
T
hof fluid inclusions in wells T-13, T-36, and T-37 (in the M-zone) show wide
variations up to 100 8C or more, and the logging temperature agrees with the
lowest homogenization temperature. The range of temperature variation at eachdepth is nearly constant and is independent of the host minerals. As discussed byTaguchi and Hayashi (1983), at the Kirishima geothermal area these thermalcharacteristics have derived from trapping of fluid in minerals throughout thecooling of the geothermal system from its peak stage.
Many T
hvalues are higher than the boiling point curve drawn from the present
water-level. As described before, many inclusions are liquid-dominant, and are notof the boiling type. This suggests that the past water-table may have been higherthan the present. The high relief (over 600 m from the wellhead of T-37 to the topof Mt Oyasu 3 km further south), deep valley, and the fluvial terrace above the
river bed in this district (Usuta et al., 1981, 1982) suggest that intensive downward
erosion caused the fall in the water-table.
As opposed to the wells in M-zone, the T
hvalues of the fluid inclusions in wells
T-21, T-29, and T-5 in C-zone vary by less than 60 8C at each depth. The logging
temperature agrees with the highest temperature, indicating that the pasttemperature may have been lower than at present, and that the temperature hasrisen and remains at its peak.
The fluid inclusion study reveals that the reservoirs in the C zone are now at
their thermal peak. In contrast, the reservoirs in the M zone have cooled fromtheir peak temperatures, and may have been a/C128ected by a drop in thegroundwater table. Thus, the Uenotai geothermal system is cooling on its margins.N. Takeno / Geothermics 29 (2000) 257–277 274
7. Conclusions
The Uenotai geothermal area is composed of a central zone of aquifer boiling
(C zone) and a surrounding liquid-dominant zone (M zone). The wells in the C
zone show excess enthalpy, and have been brought to this condition byexploitation of the original liquid-dominant reservoir. Gas equilibrium in thevapor phase is attained, and the gas composition is concordant with the vaporfraction in the reservoir and the steam quality of the discharge. A high-temperature (300 8C) high-Cl parent water is diluted before discharge from the
wells. In addition, the various geothermal waters are derived from this deepparent through boiling and vapor gain. Shallow waters derived from vapor
condensate are also present above the boiling reservoir, and drawing these waters
into wells at upper feed points modifies the discharge composition of the fluid thatis mainly derived from the deep reservoir. The marginal part of the Uenotaigeothermal system (M zone) has cooled from peak conditions, and thegroundwater level has fallen.
Acknowledgements
The author appreciates the helpful comments of Dr J. W. Hedenquist for
improving the manuscript.
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Geochemical Journal, Vol. 26, pp. 85 to 97, 1992
the Ordovician arc ophiolite,
Hayachine and Miyamori complexes, Kitakami Mountains,
Northeast Japan: isotopic ages and geochemistry
KEN SHIBATA' and KAZUHITO OZAWA2
Geological Survey of Japan, Tsukuba 3051,
and Geological Institute, University of Tokyo, Tokyo 1132, Japan
(Received January 8, 1992; Accepted March 26, 1992)
Sm-Nd, Rb-Sr and K-Ar geochronological and isotopic analyses were made on the Hayachine
Miyamori arc ophiolite in the Kitakami Mountains, Northeast Japan. The Sm-Nd isochron for four
whole-rock samples gives an age of 510 ± 70 Ma and an initial '43Nd /'44Nd ratio of 0.51229±0.00010
(ENd = + 6.4 ± 2.0). The initial ENd for individual rocks ranges from + 6.3 to + 7.0, and is lower than that
of the MORB. The age of 510 Ma is interpreted to represent the crystallization age of the Hayachine
Miyamori ophiolite. K-Ar ages of eight hornblendes range from 244 to 473 Ma, within which ages
younger than about 400 Ma are thought to be affected by Cretaceous granitic intrusions.The K-Ar horn
blende ages may indicate the times of tectonic emplacement to shallower levels.
The initial S"Sr/86Sr ratios of eight whole rock samples from the Hayachine-Miyamori complexes
show a wide variation: ranging from -4.0 to +31.9 as ESr. The nearly constant ENd and variable as, sug
gest various extent of seawater alteration for these samples. Amphibolites whose protolith may be ex
trusive or shallow intrusive rocks give higher as, than plutonic rocks, which is consistent with circulation
of sea water during the metamorphism. However, coarse-grained gabbro intruded into host peridotites
still have higher as, than the MORB. The lower ENd and higher ESr for these gabbroic rocks, which occur
in the host peridotites with arc-related geochemical characteristics, suggest that these isotopic signatures
were recorded within an arc upper mantle. This mantle signature may be explained by an involvement of
radiogenic fluid or melt derived from subducting slab, which had been altered by sea water at a mid
ocean ridge.
INTRODUCTION
Many ophiolites show unequivocal
differences from major oceanic crust and upper
mantle in many aspects (Dick and Bullen, 1984;
Dick and Fisher, 1984; Pearce et al., 1984), and
they are considered to be mostly formed in arc
related environment, such as back arc basin,
fore arc, or embryonic arc. One of the most
remarkable differences between ophiolites and
major oceanic crust-upper mantle is geochemical
characteristics of many ophiolites indicating the
similar signature of the present-day arc
magmatisms (Pearce et al., 1984). In spite of the
abundant geochemical data indicating arcsignature on extrusive or shallow intrusive rocks,
ultramafic rocks commonly do not show clear
geochemical characteristics of arc magmas. The
Miyamori ophiolitic complex is exceptional in
this sense, because its ultramafic rocks contain
abundant hydrous phases exhibiting evident arc
related geochemical characteristics (Ozawa,
1988; Ozawa and Shimizu, 1991). This indicates
that in the Miyamori complex generation
mechanism of geochemical signature for arc
magmatism can be traced back to the upper man
tle, where melting and melt segregation of
primitive magmas took place.
The Miyamori complex was transported to
the present position from the Hayachine Tec
85
86 K. Shibata and K. Ozawa
tonic Belt, where the Hayachine ophiolitic com
plex occurs. These two complexes were emplaced
in the Ordovician time along the Hayachine Tec
tonic Belt (Ozawa et al., 1988) and share the com
mon petrologic and geologic characteristics.
They can be grouped as one complex and can be
called Hayachine-Miyamori ophiolite.
In order to understand processes involved in
arc magmatisms from partial melting to volcanic
activities in the Hayachine-Miyamori ophiolite,
complete data set for trace elements and isotope
geochemistry are absolutely necessary. Ozawa
(1988) presented mineralogical data with which
qualitative characterizations of the mantle proc
esses are possible. However, those geochemical
data have not yet been obtained, because of
difficulties arising from strong alteration and
late-stage metamorphism by Cretaceous granitic
intrusions.
This paper presents Nd and Sr isotope data
of the Hayachine-Miyamori ophiolitic com
plexes along with dating by K-Ar method on
hornblende separates and Sm-Nd method on
whole rocks. Analyses were made on gabbroic
rocks because their mafic phase tends to preserve
more primary mineralogy than ultramafic rocks.
These data confirmed that the Hayachine
Miyamori ophiolite was formed in an arc en
vironment in the Ordovician time. Origin of
geochemical features characterizing arc
volcanisms are discussed on the basis of the
geochemical and mineralogical data obtained so
far.
GEOLOGIC SETTING
The Hayachine and Miyamori ophiolitic com
plexes are located in the Kitakami Mountains in
Northeast Japan (Fig. 1). The Hayachine com
plex and petrologically similar but smaller mafic
ultramafic complexes and extrusive members
probably related to the complexes define "Hayachine Tectonic Belt", which divides the
Kitakami Mountains into two geologically
distinguishable domains formed in contrasting
sedimentary environments (Onuki, 1969). The
South Kitakami is characterized by dominance Hayachine
MORoOK A"%c'y~ lltomplex +
e %3a 0 Complex HY10@ e
HIZUMT'9
0 c= °n ~IuI~T.HAY
~~ HY24
Mlyamori
Ultramafic
ComplexINE
+++
++ + +
+
MOR I+ ~
s.
N
HY23
+ o
0
4 ~G + n W
F+++ d
\++±+ HY
MLYA
45°HY15HY16
A
++K +
A+
T+\H Y20+
-I
KAMAISHI
B0!~T
39•
'KESSENNUMA
® Silurian & Devonian
® Motai Metamorphic
© Cretaceous Granitic
Uitramafic&Gabbroic
Hikami Granitic Rocks
0 20 4014 2°
I
Strata
Rocks
Rocks
Rocks
km
Fig. 1. Geologic outline of the Hayachine-Miyamori
ophiolitic complexes. Sample localities are also
shown.
of sandstone, mudstone, and limestone and has
considerable distribution of Silurian, Devonian
and possibly Ordovician sedimentary rocks
(Ehiro et al., 1988). On the contrary, the North
Kitakami is composed mainly of chert and basic
extrusive rocks of Triassic to Jurassic ages. On
the basis of this contrast and geological transi
tion from the North to South Kitakami, various
tectonic histories of the Kitakami Mountains are
proposed (Saito and Hashimoto, 1982; Osawa,
1983; Kato, 1985; Ehiro et al., 1988; Ozawa et
al., 1988).
The Hayachine and Miyamori complexes are
now separated by the Hizume-Kesennuma fault,
but they were forming one large arc ophiolitic
complex when they emplaced in the Ordovician
Ordovician arc ophiolite 87
time along the Hayachine Tectonic Belt (Ozawa
et al., 1988). The part of the large complex
(Miyamori complex) was displaced to the south
in the early Cretaceous by the left lateral strike
slip movement of the Hizume-Kesennuma fault
(Ehiro, 1977; Ozawa, 1984; Ozawa et al., 1988).
According to these tectonic reconstructions and
the common petrologic and geologic features,
the Miyamori and Hayachine complexes are
referred together as the Hayachine-Miyamori
ophiolite in this paper.
The detailed geologic and petrologic studies
of the Miyamori complex were made by Ozawa
(1983, 1988). The complex consists of ultramafic
tectonite and cumulate members that are intrud
ed by clinopyroxene hornblende gabbro,
clinopyroxene hornblendite, dolerite, and felsic
porphyritic intrusive rocks (Seki, 1952; Ozawa,
1984). The tectonite member is composed mostly
of hornblende-bearing harzburgite and dunite,
the cumulate member, mainly of hornblende
bearing dunite, plagioclase-free or -bearing
wehrlite, and olivine clinopyroxenite. They have
the typical structural and chemical characteris
tics of the basal, ultramafic portion of an
ophiolite (Ozawa, 1983).
The geology of the Hayachine complex has
not yet been well understood. There are papers
on geology of the surrounding area (Osawa,
1983; Ehiro et al., 1988), but no detailed study
on the complex itself has yet been done. A recon
naissance in the western part of the Hayachine
Tectonic Belt shows that there is rough ophiolitic
stratigraphy with substantial tectonic complica
tions. The residual peridotites occur in the cen
tral zone of the Hayachine Tectonic Belt, whose
southern margin is fringed by gabbro-diabase in
trusive complexes. To the south of this intrusive
complexes, abundant basic extrusive rocks
occur, which are most widely distributed to the
east of Hizume. These basic extrusive rocks are
covered by probably Silurian-Ordovician
sedimentary rocks in the Ohasama area (Ehiro et
al., 1988). Amphibolite and clinopyroxene horn
blende gabbros of the Hayachine complex occur
as tectonic blocks with several meter to several
tens of meter size and are in fault contact withsurrounding serpentinite or pelitic schist (Okami
and Oishi, 1983). In the Miyamori complex,
hornblende gabbro and hornblendite occur as
elongate bodies in the host peridotites. They are
interpreted to be late-intrusives (Ozawa, 1984).
PETROGRAPHY AND GEOLOGIC BACKGROUND
OF THE STUDIED SAMPLES
Sampling locality: west of Nagano Path,
Morioka City, Iwate Prefecture (39°34.0'N,
141021.0'E)
Sample no.: HY-5
Rock type: hornblende gabbro
The exposure of this gabbro is very poor, but
lineated gabbro (or amphibolite?) with similar
lithology is in fault contact with serpentinite 100
m to the NE of this locality. The gabbro shows
marked lineation and consists mainly of
brownish green to pale green hornblende
(2-0.1 mm) and saussuritized plagioclase
(2-0.1 mm). Rounded grains of clinopyroxene
are present as inclusions in or along grain bound
aries of hornblende grains. Hornblende contains
rounded plagioclase inclusions. Apatite,
0.07 0.02 mm long, is present as an aggregate.
Hornblende is very fresh with limited replace
ment by colorless amphibole along its margin.
Saussurite is composed of prehnite, albite,
grossular. Veins of prehnite, albite, and epidote
are present. Its thickness ranges from 0.5 to 0.01
mm.
Sampling locality: west of Itsutsuha, Shiwa
Town, Shiwa-gun, Iwate Prefecture (39°35.1'N,
141°17.3'E)
Sample no.: HY-9
Rock type: amphibolite
This amphibolite is in fault contact with
serpentinite on the east and is probably tectonic
block at least 5 m across. This rock consists
mainly of hornblende (1.5 0.1 mm in size) with
greenish brown to pale brown pleochroism and
saussuritized plagioclase (0.4 0.2 mm). It ex
hibits remarkable lineation, defined by linear
preferred shape orientation of hornblende,
plagioclase, and sphene. The hornblende and
88 K. Shibata and K. Ozawa
plagioclase show typical tabular equigranular tex
ture, and the texture is different from that of
epidote amphibolite occurring in a fenster of the
Miyamori complex, which is characterized by ser
rated grain boundaries of hornblende. Euhedral
sphene, 0.2 0.1 mm long, commonly occurs as
inclusions in hornblende as in the case of the
epidote amphibolite from the Miyamori com
plex. It also occurs along grain boundaries of
main constituent minerals and enclosed in
plagioclase. In any case, sphene exhibits marked
preferred shape orientation that is parallel to the
lineation. The occurrence of euhedral sphene en
closed in hornblende is a characteristics of am
phibolite or epidote amphibolite in the
Hayachine-Miyamori ophiolite. Sphene also oc
curs in hornblendite or hornblende gabbro, but
it is always secondary, replacing ilmenite or horn
blende. Hornblende is fairly fresh but tends to
become colorless or pale green along its margin
or cleavage. This alteration is generally observed
along prehnite vein (0.1 0.2 mm in thickness).
Saussurite is composed of grossular, albite,
sericite, and prehnite.
Sampling locality: southwest of Itayama, Tonan
Village, Shiwa-gun, Iwate Prefecture
(39°35.7'N, 141015.6'E)
Sample no.: HY-12
Rock type: metagabbro
The locality is within the gabbro-basic in
trusive complexes. The contact relationships
with surrounding rocks are not clear in this local
ity. The main constituent minerals are anhedral
pale green to colorless amphibole and dusty
euhedral to subhedral plagioclase. Plagioclase is
4 0.5 mm long and has fine-grained inclusions
of epidote, chlorite, amphibole, grossular, and
sericite. Pale green amphibole has brown pat
ches, suggesting that the amphibole is formed by
the replacement of brown primary hornblende.
Brown hornblende very rarely occurs as an
isolated crystal with 0.5 mm size. Plagioclase has
patches or veins of glossular or epidote. Clear
veins of albite are also observed. The minor con
stituent minerals are euhedral to subhedral
apatite (0.4.0.2 mm) and pseudomorph ofmagnetite and ilmenite aggregate (1 0.1 mm).
They are completely replaced by chlorite and
sphene. Albite veins (1 0.1 mm wide) and veins
(2.0.5 mm wide) consisting of epidote and
chlorite are present. This metagabbro is inter
preted to be a gabbro containing plagioclase,
brown hornblende, magnetite, ilmenite, and
probably substantial amounts of clinopyroxene,
which was metamorphosed at relatively low tem
perature conditions. The texture characterized
by euhedral plagioclase and anhedral mafic
minerals is different from hornblende gabbros
commonly observed in the Hayachine-Miyamori
complexes, suggesting that it is a shallow in
trusive. This is consistent with the association of
fine grained basic rocks, which are extrusive
rocks and shallow intrusives in this area.
Sampling locality: east of Kagura, Kawai
Village, Shimohei-gun, Iwate Prefecture
(39°30.9'N, 141 °38.3'E)
Sample no.: HY-15
Rock type: clinopyroxene hornblende gabbro
The sample is from gabbroic. zone of the
Kagura mass (Onuki, 1962). The gabbroic zone
is probably in fault contact with serpentinite on
the east and is gradational contact with basic in
trusive complexes on the west. This gabbro ex
hibits marked foliation and consists mainly of
brown to pale brown hornblende (5 -- 0.2 mm
long), clinopyroxene (1.5.0.2 mm long), and
saussuritized plagioclase (< 5 mm long). It also
contains magnetite and ilmenite (0.5 ~ 0.05 mm),
which are replaced by chlorite, sphene, and
green amphibole. Saussurite is composed mostly
of zoisite and contains chlorite and calcite.
Brown hornblende is replaced by green am
phibole associated with tiny sphene along its
margin and cleavage. Clinopyroxene is replaced
by pale green amphibole and chlorite along its
margin and cleavage. Veins, 0.5 0.05 mm wide,
composed of pale green hornblende, chlorite,
and calcite are present.
Sampling locality: east
Village, Shimohei-gun,
(39°30.9'N, 141 °38.3'E)of Kagura, Kawai
Iwate Prefecture
Ordovician arc ophiolite 89
Sample no.: HY-16
Rock type: clinopyroxene hornblende gabbro
The geology is the same as that of HY-15 de
scribed above. It consists mainly of brown to
pale brown hornblende (1- 0.1 mm), clinopyrox
ene (1 0.1 mm), and completely saussuritized
plagioclase. They exhibits preferred shape orien
tation, which brings about the marked foliation.
Euhedral to subhedral apatite, 0.2 0.05 mm
long, is present. Hornblende is replaced by pale
green to colorless amphibole associated with tiny
sphene along its margin and cleavage.
Clinopyroxene is replaced by chlorite, pale
brown to colorless amphibole along its margin
and cleavage. Some clinopyroxene grains are
replaced by patches of brown hornblende. As a
whole clinopyroxene is well preserved.
Saussurite is composed of zoisite, grossular, and
chlorite. Albite veins with 0.4 -- 0.02 mm
thickness are present.
Sampling locality: northwest of Osano,
Kamaishi City, Iwate Prefecture (39'16.9'N,
141 °49.3'E)
Sample no.: HY-20
Rock type: hornblende gabbro
This rock consists mainly of greenish brown
to pale brown hornblende (1 0.1 mm) and
saussuritized anhedral plagioclase (2 0.05 mm),
and exhibits weak foliation. Hornblende con
tains euhedral apatite (0.4 -- 0.05 mm) and
magnetite and ilmenite (0.2 - 0.05 mm), which
are altered into haematite. Greenish brown horn
blende rarely contains rounded clinopyroxene in
clusions 0.05 mm in diameter. Hornblende is
replaced by colorless amphibole along its margin
and cleavage and each grain has several domains
showing different extinction angle, which sug
gests plastic deformation. Crystal aggregates
composed of fine-grained colorless to pale green
acicular amphibole, chlorite, and sphene are pre
sent. They are in sharp contact with greenish
brown hornblende, suggesting that they are
pseudomorphs of clinopyroxene. Therefore, the
most of secondary amphibole listed in Table 1 is
considered to have been clinopyroxene before
the alteration. Saussurite is composed ofgrossular, albite, and chlorite, and locally cut by
albite veins. Euhedral apatite with 0.2 0.05 mm
size is present in hornblende or along grain
boundaries of the main constituent minerals.
Sampling locality: south of Kuromori, Tonan
Village, Shiwa-gun, Iwate Prefecture
(39°35.6'N, 141016.6'E)
Sample no.: HYTH3P
Rock type: amphibolite
This amphibolite occurs in the gabbro
diabase intrusive complexes. It has well devel
oped lineation and saussuritized plagioclase. It is
surrounded by much more felsic intrusive rocks
but the contact is not exposed. Hornblende is
greenish brown or brownish green to pale brown
with 1 to 0.1 mm length. Rarely large grains of
plagioclase or hornblende up to 2 mm in size are
present. They are elongate parallel to the linea
tion. Euhedral sphene with 0.05 mm size is abun
dant as observed in HY-9. Euhedral apatite with
0.03 0.01 mm size is also present. Sphene and
apatite are commonly enclosed in hornblende or
plagioclase. Hornblende is very fresh and there
is no secondary amphibole except for the margin
of grains in contact with chlorite veins. Prehnite
veins (0.1 0.05 mm in thickness) and veins com
posed of chlorite and albite (0.3 0.05 mm in
thickness) are present. Saussurite is composed of
albite, prehnite, zoisite, grossular, and sericite.
Sampling locality: northwest of Gorin Path,
Towa Town, Waga-gun, Iwate Prefecture
(39016.1'N, 141°20.9'E)
Sample no.: HY-23
Rock type: plagioclase-bearing hornblende
clinopyroxenite
This sample is from an ultramafic-mafic com
plex in the cumulate member of the Miyamori
complex. This complex might be an intrusive
into the cumulate member or melt rich pocket
surrounded by ultramafic cumulates. The com
plex consists of cortlandite, olivine hornblende
clinopyroxenite, plagioclase-bearing or -free
hornblende clinopyroxenite, clinopyroxene horn
blendite, and clinopyroxene hornblende gabbro.
They are systematically distributed from the east
90 K. Shibata and K. Ozawa
Table 1. Modal composition of the Hayachine-Miyamori ophiolitic complexes dated by K-Ar method
Sample no. Rock type cpx rcp phb rhb pl ilm apt sph** ven
Hayachine complex
HY-5
HY-9
HY-12
HY-15
HY-16
HY-20
HYTH3P
Miyamori complex
HY-23Hb gabbro
Amphibolite
Doleritic metagabbro
Cpx-hb gabbro
Cpx-hb gabbro
Hb gabbro
Amphibolite
Hb clinopyroxenitetr.
6.9
13.7
tr.
47.88.2
5.054.0
54.5
2.4
46.2
20.5
45.9
50.3
39.5 1.0
2.3
36.7
13.7
11.4
31.5
1.8
3.843.6
38.6
53.3
20.4
45.7
20.1
41.4
8.61.2
1.3
tr.
0.2tr.
0.1
0.1
tr.
0.4
tr.
0.11.2
2.81.4
3.3
6.3
3.3
3.7
2.1
3.7
-: not present, tr.: trace amount present (less than 0 .1%).
Abbreviations of mineral names are as follows; cpx: clinopyroxene, rcp: chlorite and amphibole replacing clinopyroxene,
phb: primary hornblende, rhb: secondary amphibole and chlorite replacing hornblende, pl: plagioclase (altered into
saussurite), ilm: ilmenite and magnetite (including sphene replacing ilmenite), apt: apatite, sph: sphene, ven: vein minerals
(albite, chlorite, prehnite, epidote, and calcite).*Amphibole and chlorite after clinopyroxene are included in those after primary amphibole, because of difficulty in iden
tification of primary minerals.**Primary sphene. All rocks contain trace amounts of secondary sphene.
to the west in the order listed above. This rock
has porphyritic appearance, exhibited by large
euhedral green hornblende "phenocryst", 3
cm 5 mm long, set in a medium-grained matrix
consisting mainly of euhedral to subhedral
clinopyroxene (2 0.2 mm) and interstitial
saussuritized plagioclase. Clinopyroxene is
generally replaced by patches of hornblende.
Hornblende shows pale greenish brown to pale
green pleochroism. It becomes almost colorless
with slight greenish tint where it is recrystallized
into finer grains by deformation or replaced by
chlorite. Large euhedral hornblende encloses
abundant euhedral clinopyroxene grains with the
same size as in matrix. Subhedral apatite
(0.5 0.3 mm) is present as an inclusion in large
hornblende and in matrix. Saussurite is compos
ed of zoisite, grossular, albite. Haematite,
chlorite, epidote, and sphene are present as sec
ondary phases.
Modal composition of dated samples from
the Hayachine-Miyamori complexes is given in
Table 1. Description and modal composition of
sample HY-24(AG4-2) are given in Ozawa et al.
(1988).ANALYTICAL PROCEDURES
K-Ar ages were determined on hornblendes
separated from rocks of the Hayachine
Miyamori ophiolite. The analytical method was
the same as described in Ozawa et al. (1988).
Sm-Nd and Rb-Sr analyses were made on
whole-rock samples of the Hayachine-Miyamori
ophiolite. Sm and Nd concentrations were deter
mined by isotope dilution method, whereas Rb
and Sr concentrations were determined by either
X-ray fluorescence or isotope dilution method.
The 143Nd /'44Nd and 87Sr / $6Sr ratios were deter
mined for unspiked solutions, and normalized to
146Nd / 144Nd = 0.7219 and 86Sr /81 Sr = 0.1194.
Replicate analyses of LaJolla and E & A stan
dards gave '43Nd / 144Nd and 87Sr / 86Sr ratios of
0.51184±0.00002 (2Q) and 0.70809±0.00004
(2Q), respectively. Sm-Nd age was calculated by
the least-square method of York (1966), using er
ror of 0.5% (la) for 147Sm/'44Nd ratio and errors
given in Table 3 for 143Nd / 144Nd ratio. Errors in
Sm-Nd age and initial ratio are given on 2Q
level. As the '43Nd/144Nd ratio of the LaJolla
standard in this study was lower by 0.00002 than
that of Lugmair and Carlson (1978), 0.00002 is
added for the calculation of CNd. Errors in
87Sr /16 Sr measurements were less than 0.01% at
Ordovician arc ophiolite 91
2cr mean level.
Decay constants used in age calculation
are: 40K A,8=4.962 x 10-10/y, 40K A.e=0.581 x
10-1°/y, 40K/K=0.01167 atom%, 87Rb A=1.42
X 10-11/y, 147Sm A=6.54 x 10-12/y.
RESULTS
K-Ar ages of hornblendes from five gabbroic
rocks, two amphibolites and one hornblende
clinopyroxenite are given in Table 2. Ages range
from 244 to 473 Ma, but those younger than
about 400 Ma are thought to be affected by the
Cretaceous granitic intrusions as discussed later.
Excluding them, the K-Ar ages of hornblendes
from the Hayachine complex are similar to ages
of 421-484 Ma for the Miyamori complex
(Ozawa et al., 1988), confirming that the
Miyamori complex was originally emplacedalong the "Hayachine Tectonic Belt" as a part
of the Hayachine-Miyamori ophiolite (Ozawa,
1984).
Sm-Nd isotopic data for four whole-rock
samples are given in Table 3, and are also plotted
on an isochron diagram (Fig. 2). Although four
samples are collected from a wide area in the
Hayachine and Miyamori complexes, they are
believed to be derived from a common source
material, because the terrane shares the common
characteristics in geology, petrology, and K-Ar
hornblende age (Onuki, 1962; Osawa, 1983;
Ehiro et al., 1988; Ozawa et al., 1988). The data
points define a Sm-Nd age of 510 ± 70 Ma and
an initial 143Nd/ 144Nd ratio of 0.51229±0.00010
(MSWD=0.91). The initial eNd calculated using
the parameters in Table 3 is +6.4±2.0. The ENd
calculated at 510 Ma for each sample ranges
from +6.3 to +7.0 (Table 3). Although the er
Table 2. K-Ar ages of hornblendes from the Hayachine-Miyamori ophiolitic complexes
Sample no. Rock MineralK20
(%) 40Ar rad
(10-6 MI STP/g)Atm. 40Ar
(%)Age
(Ma)
Hayachine complex
HY-5
HY-9
HY-12
HY-15
HY-16
HY-20
HYTH3PHb gabbro
Amphibolite
Metagabbro
Cpx-hb gabbro
Cpx-hb gabbro
Hb gabbro
AmphiboliteHornblende
Hornblende
Hornblende
Hornblende
Hornblende
Hornblende
Hornblende (HC1)
Hornblende
Miyamori complex
HY-23 Hb clinopyroxenite Hornblende0.129
0.137
0.134
0.121
0.106
0.294
0.293
0.135, 0.133
0.1332.06
2.39
1.13
1.24
1.22
4.91
4.86
2.04
1.96
2.3049.1
62.2
49.0
57.3
42.6
37.9
28.1
55.7
44.6
40.2Av.437±17
473 ±24
244±10
292±13
325±12
455±16
452±15
418±18
404±15
411±12
469±17
Table 3. Sm-Nd isotopic data for the Hayachine-Miyamori ophiolitic complexes
Sample no. RockSm
(ppm)Nd
(ppm)147Sm/144Nd 143Nd / 144Nd*ENd**
HY-5
HY-9
HY-23
HY-24Hb gabbro
Amphibolite
Hb clinopyroxenite
01 hornblendite0.200
2.93
1.29
2.850.521
7.78
4.20
7.540.2322
0.2278
0.1858
0.22870.51307±2
0.50304±2
0.51291±2
0.51308±6+ 6.6 ± 0.4
+6.3-L0.4
+6.4±0.4
+7.0±1.2
*Errors are 2a mean. * * Values at 510 Ma, calculated using present !43Nd /'44Nd and "7Sm / i44Nd for CHUR: 0.51264 and 0.1966 (Wasserburg et al.,
1981).
92 K. Shibata and K. Ozawa
0.5135
0.5130
Z 4
a z
0
0.5125
0.5120T= 510 ±70 Ma
I = 0.51229 ± 0.00010
e=+6.4±2.0
0.10 0.15 0.20 0.25 0.30
147Sm/144Nd
Fig. 2. Sm-Nd isochron diagram for whole-rock
samples from the Hayachine-Miyamori ophiolite.500
400
Q Y
0)
c
300
c L 0
2
O
0
0
This study
O Ozawa et al. (1988)
O
ror in age is relatively large owing to a limited
range in 147Sm/ 144Nd ratio, the age of 510 Ma
predates the K-Ar hornblende ages, and is inter
preted to be the crystallization age of the
Hayachine-Miyamori ophiolite. The ENd of +6.4
is slightly lower than that of the MORB even cor
rected for the age of 510 Ma and is within the
range of island arc tholeiites.
Rb-Sr isotopic data are given in Table 4. In
itial 87Sr/86Sr ratios calculated for the Sm-Nd
age of 510 Ma show a wide variation from
0.70361 to 0.70614, corresponding to esr of -4.0
to + 31.9. The pattern of limited ENd and variable
Esr is similar to that of the Bay of Islands
(Jacobsen and Wasserburg, 1979) and Samail
(McCulloch et al., 1980) ophiolites, suggesting
the effect of seawater alteration. This problem
will be discussed later.
DISCUSSION
Evaluation of K-Ar ages
The investigated rocks show a large variation
in degree of secondary alteration or low-tempera
ture metamorphism, although they are supposed
to have been equilibrated at higher amphibolite
facies as indicated by their primary assemblage
of brown hornblende, clinopyroxene, and prob
ably calcic plagioclase. Total modal percentage
of the secondary mafic minerals ranges from 2 to
43%, in which secondary amphibole occupies
the most of them (Table 1). The obtained K-Ar
ages of hornblende show significantly wide range0 20 40 60 80 100
% of secondary mafic minerals in total mafics
Fig. 3. Percentage of secondary mafic minerals in
total mafics against hornblende K-Ar ages.
from 244 to 473 Ma. In order to evaluate the
effects of these secondary processes, K-Ar ages
are plotted against percentage of secondary
mafic minerals in total mafic minerals in each
rock as an indicator of degree of alteration
and/or metamorphism (Fig. 3).
K-Ar ages exhibit negative correlation with
the percentage of secondary mafic minerals as a
whole, which is consistent with the data on the
Miyamori complex (Ozawa et al., 1988). This cor
relation clearly suggests that K-Ar ages are reset
by a metamorphism that took place in ages
younger than 250 Ma in those rocks with abun
dant secondary mafic minerals. Metagabbro
HY-12 with highest percentage of secondary am
phibole indicates 244 Ma K-Ar age, of which
younger age may be attributable to one of small
Cretaceous granitic intrusives commonly occurr
ing in the western part of the Hayachine Tec
tonic Belt. The locality is 2 km to the south of
outcrops of metaperidotite with metamorphic
assemblage of talc, tremolite, and olivine. This
assemblage indicates metamorphic temperature
higher than 700°C. Amphibolite HYTH3P,
sampled 1.4 km to the east of HY-12 locality
yields 411 Ma, and amphibolite HY-9 sampled
1.5 km to the southeast of HYTH3P locality
yields 473 Ma. Very fresh peridotites with small
amount of lizardite and chrysotile as serpentine
minerals are cropped out near HY-9. The abun
dance of secondary mafic phase decreases and
K-Ar age increases from HY 12 to HY-9 through
Ordovician arc ophiolite 93
HYTH3P, suggesting that thermal effect of the
Cretaceous granitic intrusive decreases from the
west to the east in this region, and that 473 Ma is
a reasonable estimation of cooling age for the
Hayachine ophiolitic complex. Two samples
from the Kagura body indicating K-Ar ages of
292 and 325 Ma contain abundant secondary
mafic minerals and they are eventually sampled
closer to the Tono granitic intrusion. Peridotites
occurring near the sampled gabbros -have
metamorphic mineral assemblage of antigorite
and olivine, which is consistent with observed
younger ages for the gabbros as in the case of the
Miyamori complex (Ozawa et al., 1988).
Gabbro HY-20 sampled from a locality to
the west of Kamaishi is exceptional, indicating
452 and 455 Ma K-Ar ages, although it contains
considerable amounts of secondary amphibole.
Petrographic observation demonstrates that sec
ondary amphibole was mostly derived from the
replacement of clinopyroxene. The granite out
crop nearest to the gabbro locality is 10 km to
the south. Serpentinized ultramafic rock occurr
ing within a few hundred meter to the south of
this locality is characterized by static replace
ment of olivine by lizardite and/or chrysotile
assemblage, and no antigorite is present. These
facts strongly suggest that contact metamor
phism by Cretaceous granitic intrusives did not
cause replacement of those clinopyroxene in this
region. The alteration of clinopyroxene could
have been related to the emplacement stage tak
ing place at 450 Ma.
Excepting the samples containing abundant
secondary amphibole and occurring close to the
granitic intrusions, obtained K-Ar ages range
from 437 to 473 Ma, which coincide with the K
Ar ages (421 484 Ma) obtained for the
Miyamori complex (Ozawa et al., 1988). This
coincidence strongly supports the idea that the
Miyamori complex is a displaced portion of the
Hayachine ophiolitic complex, proposed by
Ozawa (1984).
Sm-Nd age and its relationship to K-Ar age
A Sm-Nd whole-rock isochron age of
plutonic rocks is generally interpreted to indicatea crystallization age, because the Sm-Nd system
is less susceptible to alteration and metamor
phism than the Rb-Sr and K-Ar systems.
Therefore, the Sm-Nd whole-rock age of
510 ± 70 Ma is interpreted to be the time of
crystallization of gabbroic rocks of the
Hayachine-Miyamori ophiolite in the mantle or
in the lower crust.
K-Ar ages for hornblendes from the
Hayachine-Miyamori ophiolite range from 421
to 484 Ma, excluding younger ages caused by
granitic intrusions. The closure temperature of
Ar diffusion in hornblende is estimated to be
about 500°C. Thus K-Ar ages of 421-484 Ma
for hornblendes indicate the times when the
Hayachine-Miyamori ophiolite cooled to about
500°C after the formation. These ages could in
dicate either the times of tectonic emplacement
to shallower levels or late stage tectonic distur
bance. The time difference between the Sm-Nd
age and the oldest hornblende age: 30 Ma, may
therefore be the period between the crystalliza
tion and tectonic emplacement of the
Hayachine-Miyamori ophiolite. However, a
large error in the Sm-Nd age makes it difficult to
further discuss this time difference. It appears
that the tectonic emplacement occurred soon
after the crystallization at least in some part of
the Hayachine-Miyamori ophiolite.
Nd and Sr isotopic systematics and origin of the
Hayachine-Miyamori complexes
The initial isotopic compositions of Sr
calculated for the 510 Ma Sm-Nd age show very
wide variation ranging from 4.0 to + 31.9 for
Esr, and from 0.70361 to 0.70614 for initial
87Sr/86Sr ratio (Table 4). This wide variation is
true even if highly altered gabbros are excluded.
There is no over-all systematic correlation be
tween initial 87Sr/86Sr ratio and abundance of sec
ondary mafic minerals. For highly altered gab
bros, which give younger K-Ar ages, the ratio
tends to decrease as secondary mafic minerals in
crease, but for fresh samples it shows wide varia
tion without any correlation with the abundance
of secondary mafic minerals. This fact clearly in
dicates that late-stage metamorphism or altera
94 K. Shibata and K. Ozawa
Table 4. Rb-Sr isotopic data for the Hayachine-Miyamori ophiolitic complexes
Sample no. RockRb
(ppm). Sr
(ppm)87Rb / 6Sr 87Sr/ 86Sr(87Sr/86Sr)* *ssI***
Hayachine complex
HY-5
HY-9
HY-12
HY-15
HY-20
HYTH3P
Miyamori comple
HY-23
HY-24Hb gabbro
Amphibolite
Metagabbro
Cpx-hb gabbro
Hb gabbro
Amphibolite
x
Hb clinopyroxenite
01 hornblendite22.77*
3.4
7.5
0.868*
3.9
1.1
0.817*
0.103*342.8*
233
312
405
110
128
55.6
65.60.1924
0.042
0.070
0.0062
0.079
0.025
0.043
0.00450.70500
0.70644
0.70467
0.70504
0.70488
0.70454
0.70456
0.705070.70361
0.70614
0.70416
0.70500
0.70431
0.70439
0.70425
0.70504 -4.0
+31.9
+3.9
+15.7
+5.9
+6.7
+5.1
+16.3
*Isotope dilution; others X -ray fluorescence.
* * Values at 510 Ma. *** Values at 510 Ma, calculated using present 87Srl86Sr and 87Rb/86Sr for UR: 0.7045 and 0.0839 (DePaolo and Wasserburg,
1976).
10
8
6
v
Wz 4
2
0
-2
-20 -10 0 10 20 30 40
ESr
Fig. 4. es,-ENd diagram for the Hayachine
Miyamori ophiolite. Closed circles represent samples
of which ENd values are assumed to be + 6.4. The
horizontal line is a part of mixing curve between the
uncontaminated rock and the Ordovician seawater.l1 O,
eO
,., . O
i i
tion does not account for the wide variation of in
itial S7Sr / 86Sr ratio.
In contrast to the wide variation of Esr, eNd is
almost constant ranging from 6.3 to 7.0 giving
rise to a nearly horizontal trend on an ES,, CNd
diagram (Fig. 4). This trend coincides with that
reported from the Bay of Islands Complex,
which is attributed to the seawater alteration
(Jacobsen and Wasserburg, 1979). The sample
with highest as, is amphibolite (+ 31.9), which is
supposed to be metamorphosed basalt, diabase
dike, or gabbro. Another amphibolite HYTH3P
also indicates the second highest Esr value of
+6.7 if gabbros with high modal abundance of
secondary mafic phases and olivine hornblenditeHY-24 are excluded. HY-24, which is peculiar in
the presence of olivine, orthopyroxene, and
spinel, absence of plagioclase, higher Mg# of
mafic phases, and higher concentration of K20
in hornblende (0.236 wt%), has the highest ENd
of 7.2. Its asr is also relatively high and indicates
+ 16.3. The other two samples are coarse-grain
ed hornblende gabbro and hornblende
clinopyroxenite which preserve magmatic tex
tures under slow cooling condition. This correla
tion between rock type and asr is consistent with
the idea that the wide 87Sr/86Sr ratio is originated
from various degree of seawater interaction with
a cooling magma system; extrusive or shallow in
trusive rocks were highly interacted with
seawater and recrystallized into amphibolite
with strong lineation, whereas plutonic rocks
were not. This is similar to the Bay of Island com
plex, in which metagabbro, trondhjemite, and
sheeted dikes give higher Es,,, whereas fresh gab
bros and peridotites give lower Esr (Jacobsen and
Wasserburg, 1979).
The elevated initial 87Sr/86Sr for the gabbroic
rocks could be explained by seawater circula
tion, but these hornblende gabbros and horn
blendites are present as intrusives in ultramafic
rocks, suggesting that seawater circulation dur
ing the formation of shallower crust may not be
responsible for the origin of their high as,. Even
if amphibolites are neglected, the data are plot
Ordovician arc ophiolite 95
ted above the Mantle Array, the fact of which
have been reported from some island arc
volcanic rocks (Hawkesworth et al., 1977;
Hawkesworth et al., 1979; Whitford et al., 1981).
The ENd values are lower than those of MORB
and are within the range of some island arc
volcanic rocks (Nohda and Wasserburg, 1981;
McCulloch and Perfit, 1981).
The fact that enrichment in radiogenic Sr can
be observed in the mantle of the Hayachine
Miyamori ophiolite is very important, because
any shallow processes may not account for the
high 87Sr / 86Sr ratio of coarse-grained basic intru
sion in the mantle peridotites, which preserve
high pressure-temperature primary assemblages.
There is no evidence for the participation of
crustal contamination in the formation of the
ophiolite. The ENd Esr systematics is a mantle
geochemical signature, although mantle
peridotites or pyroxenites have to be analyzed
for the confirmation, which is an important
future work.
On the basis of mineralogic data of am
phibole in peridotites, Ozawa (1988, 1990)
argued that the Miyamori complex was formed
in an arc environment. Yoshida et al. (1990)
reported that extrusive rocks from the
Hayachine Tectonic Belt show geochemical char
acteristics of arc magmatism. Olivine-phyric
basalts occurring as volcanic breccia in the
southwestern part of the Hayachine ophiolitic
complex is highly vesiculated, even if it is quite
primitive. The high vesicularity suggests high
water content unless they erupted on very
shallow ocean basin (Dick, 1980). Spine)
phenocrysts and inclusions in altered olivine
phenocrysts of those basalts exhibit Cr# as high
as that in chromian spine) peridotites of the
Miyamori complex (Ozawa, 1988). As shown in
Fig. 5, the spine) compositions are partly within
the range of boninite and high-Mg andesite
(Dick and Bullen, 1984). Some of them are
within MORB area, but they tend to show higher
ferric and ferrous contents, suggesting
crystallization from evolved Welt. All these data
suggest that the Hayachine-Miyamori ophiolite
represents a section of arc crust and upper man1.0
0.8
0.6
U
U 0.4
0.2
0.0Hayachine spine) in basalt
0 IBoninite and
high-Mg andesite
ii0`'
r~0
80102~~o
O80102015C-1
801020150-2
801021050
80101015E-1
1
Fig. 5. Composition
Hayachine Tectonic
phenocrysts of tiny in
which are completely p
localities are Otsube to the east of Hizume.0.8 0.6 0.4 0.2 0
Mg/(Mg+Fe 2+)
of spinet in basalts from the
Belt. Spinet grains occur as
clusions in olivine phenocrysts,
altered into chlorite. Sam ding
tle. The high initial 87Sr/86Sr ratio of the
Hayachine-Miyamori ophiolite, therefore,
should be considered in this context.
Gabbroic rocks in the Miyamori complex are
distributed in residual peridotites as elongate
bodies (Ozawa, 1984). As Seki (1952)
demonstrated, at least some of them are dikes in
truded into the host peridotites. These dikes,
whose contact relationships with its host can be
observed, are small, usually less than a meter in
thickness, and mostly hronblendite or
clinopyroxene hornblendite. They do not con
tain plagioclase at all. A large gabbroic body oc
curring to the north of Miyamori has thick
selvage of hornblendite on the both sides
(Ozawa, 1983). The occurrence of hornblendite
between hornblende gabbro and peridotite sug
gests reaction between hydrous basic melt and
host peridotites. The cumulate member of the
Miyamori complex represents late stage
cumulate body intruded into the tectonite
member (Ozawa, 1983). The common occur
96 K.Shibata and K. Ozawa
rence of late intrusions often associated with
reaction zone may be highly expected in the arc
upper mantle, where magma may be supplied
repeatedly. The observed high Es, and low eNd for
these intrusives may represent isotopic character
istics of the arc magmatisms that took place in
the Hayachine-Miyamori ophiolite. This
isotopic signature may be attributable to the in
volvement of altered old oceanic crust or ter
rigenous sediments subducted beneath the
Hayachine-Miyamori arc in the Ordovician
time.
Acknowledgments-We wish to thank Dr. T. Tanaka
of Nagoya University for assistance and discussion in
Sm-Nd dating, and Mr. S. Uchiumi of the Geological
Survey of Japan for assistance in K-Ar dating.
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Shibata & Ozawa 1992 Ordovician arc ophiolite, the Hayachine and Miyamori complexes, Kitakami Mountains, Northeast Japan, isotopic ages and geochemistry.txt
|
The Island Arc (1992) 1, 2-15
Review Article
Crustal structure and origin of the northeast Japan arc
KOJI MINOURA' AND AKIRA HASEGAWA~
'Institute of Geology and Paleontology and 'Observation Center for Prediction of Earthquakes and Volcanic Eruptions,
Faculty of Science, Tohoku University, Sendai 980, Japan.
Abstract Northeast Japan is a typical island arc region and its topographic arrangement reflects the
geophysical characteristics of the island arc system. However, the structural style of the arc is very
complicated and varied due to the repeated superposing of faults and folds on to earlier structures.
Geotectonic events that involved creation of the fundamental framework of the island arc crust
occurred in east Asia in the Late Jurassic to Early Cretaceous and were probably induced by accretion and
collision tectonics. The fragmentation and subsequent displacement of the crust took place during the
Early Neogene in response to the terrane collision and the change in oceanic plate motion, leading to the
opening of the Japan Sea. Huge amounts of volcano-sedimentary rocks buried the tilted fault blocks of
pre-Tertiary basement with the development of the island arc.
Key words: calc-alkalic magmatism, collision suture, gravity anomaly, island arc, subduction complex.
INTRODUCTION
Northeast Japan is comprised of three gross morphostruc-
tural provinces: inner, central and outer zones, which are
arranged in a sub-parallel manner from west to east. Each
zone runs parallel to the Japan trench. The Ou backbone
range is the greatest mountain chain in the Japanese Islands
and lies north to south for over 800 km in the central zone.
The northward extension of the range, interrupted by the
Tsugaru Straits once, is succeeded by mountainous areas in
southwest Hokkaido and then dips into the Japan Sea
further north. The backbone range is topped by active or
recently active volcanoes that lie at almost regular intervals
forming a long volcanic chain parallel to the Japan trench.
The inner zone, facing the Japan Sea on the western
border, is marked by a discontinuous row of hilly mountain
masses. Volcanic activities are often recognized in the
marginal part of each mountain mass. A pair of plateaus
exhibiting an echelon arrangement is present in the outer
zone in a spindle shape with axes extending north and
south. Each plateau, margined by a large river (the Abu-
kuma or Kitakami River) on the west and a bold shore
(Sanriku coast) on the east, has received considerable
peneplanation.
Each topographic elevation of the arc is separated by a train
of narrow valleys. Anticlinal arches with north-trending
axes form the mountainous region, and the corresponding
characteristics are found in the valleys where synclinal
troughs are buried in thick fluvial deposits (Oka 1986).
Such a close correlation of topographic arrangement with
general structural trends can be recognized in southwest
Hokkaido and the extreme north of Honshu. The regional
topography of northeast Japan is schematically illustrated in
Fig. 1, showing a zonal arrangement of morphostructural
divisions.
The outer zone represented by the area of the Abukuma
and Kitakami massifs is characterized by a wide exposure of
the Jurassic accretionary complex, a pre-Cretaceous
Accepted for publication 28 March 1992. shallow marine sequence and Lower Cretaceous plutonic
rocks (Minoura 1990). Neogene volcano-sedimentary rocks
are the main constituents of the central and inner zones,
where highly sheared and mildly metamorphosed rocks are
exposed and unconformably overlain by the Neogene
cover. They bear a striking resemblance in lithology to the
subduction complex of the Kitakami massif which contains
allochthonous blocks of cherts and limestones that yield
pre-Jurassic fossils (Minoura 1985, 1990). Radiometric dat-
ing, mainly by the K-Ar method, indicates that the
plutonic intrusions emplaced in pre-Tertiary rocks are
mostly Early Cretaceous, ranging from 120 to 110 Ma
(Kawano & Ueda 1967; Shibata et d. 1978). It is suggested,
therefore, that the Jurassic subduction complex, together
with Lower Cretaceous granitic rocks, underlies the
Neogene sequence and forms the basement of the arc.
It is now generally accepted that northern Honshu is a
typical region of the island arc. Previous papers explain that
a westward-dipping sharp seismic zone, trenchward lower-
ing heat flow, and a pair of gravity high and low are
exposed to geophysical east-west cross sectional views (e. g.
Yoshii 1977, 1979). Island arc volcanism is proved by
Quaternary volcanoes. The northeast Japan arc is under
northwest-southeast compressional stress as estimated from
focal mechanisms of shallow earthquakes and horizontal
strains of the crust (e.g. Sato et al. 1980; Chida & Sato
1987). Although the topographic arrangement of the arc,
reflecting present geophysical settings, was established in
the Early Pleistocene (Minoura & Kosuga 1989), the struc-
tural style of the crust may be ascribed to different tectonic
controls of older ages.
Recent progress in investigations of Phanerozoic rocks
reveals that the northeast Japan arc has experienced a
complex history of geotectonic evolution since the Or-
dovician (Minoura 1990). Although stratigraphic analyses
of the pre-Neogene system have been the subject of geo-
logical study during the past three decades, a general
interpretation of the tectonic significance of the sedimentary
Structure and origin of the northeast Japan arc 3
basins in northeast Japan has not been possible until now.
Physical information about crustal structure is essential
to an overall estimate of the island arc situation. In assessing
the regional tectonic significance, an over-simplification of
the large-scale structural pattern is effective. In this paper,
the geological and geophysical results on the northeast
Japan arc are interpreted synthetically in order to present a
new view of the tectonic development of the arc and the
opening of the Japan Sea.
GEOPHYSICAL SETTING OF THE NORTHEAST JAPAN ARC
Northeast Japan has been regarded as a model field of
geophysical research because of its tectonic environment,
which is typical of an island arc system. Many valuable
results have been obtained from the study of this arc (e.g.
Hasegawa et al. 1978; Yoshii 6r Asano 1972), the most
valuable being the discovery of the double-planed structure
of the deep seismic zone (Umino & Hasegawa 1975). In this
section, important geophysical data are presented about the
island arc crust. These data will be referred to in a discus-
sion of the structural evolution of the northeast Japan arc.
P-WAVE VELOCITY AND MOHO DISCONTINUITY
Figure 2 illustrates the focal depth distribution of earth-
quakes in northern Honshu, projected on the vertical
section across line a-b. The double-planed structure of the
westward dipping seismic zone is discernible and the upper
plane is particularly well-defined. Almost all of the shallow
earthquakes occur in a layer that is shallower than the
Conrad discontinuity, which implies that the lower crust
has no bearing on earthquakes (Takagi et al. 1977). Figure
3 is a map of northeast Japan showing the contours of the
hypocentral depths to the upper and lower seismic planes of
the double-planed deep seismic zone (Hasegawa et al.
1983). Figure 4a, b are maps of northern Honshu showing
the depth distribution (km) of the Moho and Conrad
discontinuities, respectively (Zhao et al. 1990). Figure 4c
shows a variation in the P-wave seismic velocity (km/s) for
the upper crust (Zhao 1988). A P-wave seismic structure of
the crust and upper mantle projected on the profile across
line A-B (Fig. 4e) is illustrated in Fig. 4d.
Takagi et al. (1977) considered that the upper and lower
crusts of the northeast Japan arc are mainly composed of
felsic (granitic) and mafic (basaltic) rocks, respectively. A
geophysical profile of northern Honshu, by Yoshii (1979),
clearly shows that the upper crust decreases in thickness
just beneath the Sanriku coast and lower-crust materials
come up to a depth of 10 km. The emplacement of mafic
bodies into the felsic layer seems to be a remarkable
phenomenon that is closely related to the origin of the
island arc crust.
MAGNETIC AND GRAVITY ANOMALIES
Positive magnetic and free-air gravity anomaly maps modi-
fied from Segawa & Furuta (1978) and Tomoda (1973) are
shown in Fig. 5. The overall view of the free-air gravity
Fig. 1 Topographic division of the northeast Japan arc. A: Pre-Tertiary
basement; 6: Hilly mountain masses; C: Quaternary volcanoes. H-K: Hakodate-
Kuromatsunai depression (Oka 1986); N-M: Noheji-Mutsu depression. DE:
Dewa; TA: Taiheizan; SH: Shirakami; TU: Tsugaru; 0s: Oshima; KA: Karneyama.
14401738, 1992, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1992.tb00053.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
4 K. Minoura and A. Hasegawa
b
Fig. 2 Focal depth distribution of earthquakes, located by the seismic network of Tohoku University, from January 1981 to July 1988, projected on the vertical section
across line a-A-B-b on the inserted location map. Focal mechanisms of earthquakes beneath northern Honshu are schematically represented by solid arrows. An open
arrow with letters J.T indicates the position of the Japan trench. MI -M5: Seismic magnitude.
140' 144'
4O
40'
-
- 36'
I I I
138" 142' 146' Fig. 3 Contours of hypocentral depths to the
upper and lower seismic planes. Open tri-
angles denote the locations of active vol-
canoes. After Hasegawaet a/. (1983)
14401738, 1992, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1992.tb00053.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Structure and origin of the northeast japan arc 5
-
ao- (C) 140. 142.
6.2 - '
Depth distribution of the Moho Depth distribution of the Conrad
discontinuity (km) discontinuity (krn) (km/s) P-wave velocity in the upper crust
5.9 6.0 5.9 5.8 5.9 6.0 6.1 -
km P-wave velocity structure along profile A-B (km/s)
Fig.4 Maps showing depth distribution of the Moho (a) and Conrad (b) discontinuities and P-wave velocity in the upper crust (c); (d) shows P-wave seismic structl;re
across line A-B on E. Modified from Zhao (1988) and Zhaoet a/. (1990).
Fig.5 Magnetic high (a) and free-air gravity
(b) anomaly patterns of the northeast Japan
arc. The stippling designates the areas of
positive magnetic anomalies. Dotted (L) and
stippled (H) areas in the free-air gravity anomaly
map denote gravity lows and highs, respectively.
Modified from Segawa and Furuta (1978).
14401738, 1992, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1992.tb00053.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
6 K. MinouraandA. Husegawa
map markedly demonstrates the existence of a high gravity
anomaly zone along the Sanriku coast and the correspond-
ing characteristics are found in the magnetic anomaly
patterns.
A pair of gravity high (H) and low (L) in Fig. 5b suggest
that the northeast Japan arc is a typical island arc region.
The positive magnetic belt as well as high gravity lineation
seems to be a reflection of a latent body of higher density
rocks in the upper crust. The possible existence of dense
masses in the shallow layer of the island arc crust has
already been inferred from the seismic evidence. Segawa
and Furuta (1978) suggested that a dyke-like mafic body
rises to the depth of about 5 km under the Sanriku coast.
The Japan trench is in turn marked by an anomalous zone
of low gravity caused by a thick wedge of lower-density
materials resulting from the down-going slab at the trench
margin.
THERMAL STRUCTURE OF THE CRUST
Figure 6a is a simplified map of heat flow distribution
(modified from Yoshii (1979)), and Fig. 6b shows contours,
in km, of Curie point temperature depth (CPTD; after
Okubo (1984)). Heat flow increases at the arc-back-arc
region and decreases at the trenches and ocean, showing a
clear relationship between the island arc and the adjoining
ocean floor. The CPTD surface in Fig. 6b presents certain
anomalies: active volcanic activity accompanies a 3-4 km
rise in CPTD and thick sediments depress the CPTD in
central Hokkaido and off the Sanriku coast as well as in the
Kitakami area. The out-of-phase behaviour with a depress-
ing trend in the CPTD surface in Kitakami is recognized
just beneath the serpentine melange zone that separates the
Kitakami massif into northern and southern terranes (Mi-
noura 1983). This suggests a localized rising of heated
materials through crushed zones. A sharp thermal gradient
along the coast of the Japan Sea seems to result from high
heat flow in the back-arc areas.
REGIONAL GEOLOGY OF NORTHEAST JAPAN
A geological sketch map of northeast Japan is presented in
Fig. 7. As previously stated, the northeast Japan arc is mostly underlain by pre-Tertiary rocks and Upper Cenozoic
volcano-sedimentary covers. The exposures of pre-Tertiary
formations are so limited that the geotectonic details re-
main mysterious. This section presents a recent view of the
tectonic development of the pre-Tertiary systems presented
by one of the authors (Minoura 1985, 1990). The opening
of the Japan Sea seems a major event responsible for the
creation of the present crustal framework of northeast
Japan. Thus the Neogene development, showing a complex
phase of tectonic evolution, will be interpreted in the last
phase, in relation to the origin of the Japan Sea.
GEOLOGY OF PALAEOZOIC AND MESOZOIC TERRANES
Pre-Tertiary rocks of the northeast Japan arc are most
widely exposed in the Kitakami and Abukuma areas, where
sedimentary terranes can be classified into two categories:
Southern Kitakami and Abukuma terranes; and Northern
Kitakami terrane. Highly sheared sedimentary rocks patch-
ily distributed in the central and inner zones are possible
candidates for pre-Tertiary strata comparable with those of
the Northern Kitakami terrane.
Southern Kitakami and Abukuma terranes
The oldest rocks found in northeast Japan are Ordovician
ultramafic bodies exposed in the Southern Kitakami area
(Ozawa et al. 1988). Metamorphic rocks and granitoids
suggest that older ages have limited exposure there. The
stratigraphic succession of the Kitakami and Abukuma
terranes can be divided into two units: a Siluro-Carbon-
iferous active margin sequence and a Permian to Early
Cretaceous passive margin sequence (Minoura 1985). Sedi-
mentary history of the Southern Kitakami terrane is
schematically illustrated in Fig. 8.
The lower sequence includes huge volumes of rhyolitic to
dacitic rocks in calc-alkalic series. The upper sequence,
reaching up to 8000 m thick, is made up of shallow marine
clastic rocks and limestones, with little volcanic materials,
followed by the accumulation of arkosic sandstones with
interbeds of fossiliferous mudstones yielding abundant ammo-
nites of Oxfordian to Berriasian age (Takahashi 1969,
1973).
Fig.6 (a) Simplified map of heat flow
distribution in and around the Japanese
Islands, showing increased heat flow at
the volcanic arc and back-arc regions.
Modified from Yoshii (1979). Numbers
1 .O-2.5 show HFU values. (b) Topography
of Curie point temperature depth surface
in the northeast Japan arc. At depths
below the surface spontaneous magnetiza-
tion of ferromagnetic minerals is lost. After
Okubo (1984). Contour interval is 2 km.
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Structure and origin of the northeast Japan arc 7
The composition of conglomerates above the unconfor-
mities changes gradually from a polymictic to monomictic
type through the Permian to Jurassic. This suggests a
systematic lithofacies change of sediment provenance (Fig.
8). During the Permian and Triassic volcano-sedimentary
covers of island arc crusts were denuded and in the Jurassic
plutonic rocks were widely exposed in the sediment source
areas. Volcaniclastic rocks of bimodal characteristics, con-
taining intercalated mudstone beds rich in radiolarian fossils
of Late Valanginian age (Taketani 1987), unconformably
overlie the upper sequence.
A large structural event that occurred in the Middle to
Late Carboniferous, changed the tectonic setting of the
terrane from an island arc into a matured continental
margin (Fig. 8). Systematic changes in biofacies (Saito &
Hashimoto 1982) and palaeolatitudes (Fujiwara 1968) over
time suggest a northward drift of the terranes during the
Permian to Jurassic and a juxtaposition at the eastern
margin of the Asian continent in the Early Cretaceous
(Minoura, 1985). It is suspected that the allochthonous
land-masses originated in the lost continent of Gondwana.
Northern Kitakarni terrane
The Northern Kitakami terrane comprises two different
Jurassic subduction complexes arranged in zonal structures:
zone I composed of highly sheared and mildly metamor-
phosed mdange blocks characterized by the incorporation
of exotic mafic rocks; and zone I1 composed of jumbled
masses intermixed with andesitic volcanic rocks. The dif-
ference between these two zones can also be recognized in
the composition of sandstones; zone I is volcanic and zone
I1 is plutonic in provenance (Minoura 1990). Carbon-
iferous to Triassic fossils are found in cherts and limestones
which were emplaced, in a mixed manner, into a turbiditic
matrix or slope-failure deposit of the zones in the Middle to
Late Jurassic. There is a polarity in the younging trend of
melange blocks, suggesting the west to east growth of each
subduction complex (Minoura 1990).
Zone I1 of the Northern Kitakami terrane seems to be an
exotic fragment of the Triassic to Jurassic unmatured island
arc developed in a different tectonic province (Minoura
1985). Field observations reveal that zone I is in fault
contact with zone 11. It can be interpreted, therefore, that
the boundary dividing the Northern Kitakami terrane into
two tectonic zones is a collision suture (Fig. 7).
On the top of zone I1 there are large thrust sheets that
consist of a Triassic to Jurassic andesite-limestone-chert
sequence and a chert-shale-turbidite sequence. Valangini-
an to Hauterivian outer-shelf sediments with andesitic to
rhyolitic lavas and tuff breccias of calc-alkalic series uncon-
formably overlie the latter sequence (Minoura & Tsushima
1984). Deposition of Kuroko-type ores is recognized in the
fine clastic rocks (Yamaoka 1983). The facies characteris-
tics as well as differences in the composition of sedimentary
--t
Fig. 7 Geologic map of northeast Japan showing principal stratigraphic and
structural components. (a) Silurian to Jurassic sedimentary rocks; (b) Jurassic
subduction complex including exotic mafic to ultramafic rocks (B1 : Zone I) and
andesitic volcanic rocks (B2: Zone 11); (C) Lower Cretaceous plutonic rocks; (D)
Lower Cretaceous volcanic rocks of bimodal characteristics (D,) including
Kuroko-type ore deposits (D2). (E) Neogene volcano-sedimentary rocks; 1 : Taro
Fault; 2: Kuzurnaki Fault; 3: Hayachine Fault; 4: Hizume-Kesennuma Fault;
5: Hitokabe-lriya Fault; 6: Futaba Fault; 7: Hatagawa Fault; 8: Tanakura Fault. ---__ 7- - - - -
B
61 82
Yo bOkm 100 A 8-
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8 K. Minoura and A. Hasegawa
rocks markedly distinguish the Northern Kitakami terrane
from the Southern Kitakami and Abukuma terranes (Katada
& Ono 1978).
CENOZOIC VOLCANO-SEDIMENTARY COVER
Latest Palaeogene time was the most important period in
establishing the fundamental framework of the island arc
crust. During the Late Oligocene to Early Miocene, inten-
sive and extensive volcanism of bimodal characteristics
took place in the major part of northeast Japan. Huge
volumes of volcanic materials accumulated over the pre-
Tertiary rocks (Yoshida 1975). The great volcanic event
was followed by back-arc basin subsidence that continued
throughout the Middle Miocene (Susaki & Minoura 1992)
under the tectonic control of east-west extensional stress
(Otsuki 1990). It is possible that the basin subsidence
resulted from the displacement of faulted landmasses.
Figure 9 shows palaeogeographic maps of southwest Hok-
kaido and the extreme north of Honshu at various times
since 18 Ma. These maps were reconstructed from many
stratigraphic data (Susaki & Minoura 1992). Tectonic
processes estimated from Miocene stress data for the arc
coincide with the basin development and show a consistent
pattern of extensional stress that is concordant with basin
subsidence.
In a parallel fault network, a series of extensional tec-
tonic processes led to alternate zones of raised and de-
pressed fault blocks on the eastern margin of the Eurasian
plate. This resulted in the development of submarine swells
and furrows (Fig. 9). Swells lying north and south were
separated by trains of furrow that subsided a few hundred
metres below the average sea floor (Susaki & Minoura
1992). The principal part of the furrows was relatively
starved of sediments and filled only with siliceous plank-
tonic shells and small amounts of fine clastic materials.
Basaltic volcanic events successively occurred on the furrow
axis and were followed by andesite to rhyolite volcanic activities. The inter-furrows and furrow margins, in turn,
became a major location of sediment accumulation and
received intermittently voluminous pyroclastic flows from
submarine volcanoes. The deposition of black-ores is be-
lieved to have occurred as a result of the underwater
volcanism that took place in the furrows in the Middle
Miocene (Kitazato 1983).
Figure 10 illustrates the distribution of volcanic products
of the Late Oligocene to Early Miocene, in relation to the
principal faults (Minoura 1989). The occurrence of vol-
canic rocks along the faults seems to simply suggest that the
intensive volcanism was caused by magmatism associated
with the fragmentation of the island arc crust. The opening
of the Japan Sea may have been a consequence of the
oceanward rafting of the fragmented crusts.
Cenozoic igneous activities, leading to a voluminous
accumulation of volcaniclastic materials, are classified into
two stages: the Late Oligocene to Early Miocene; and the
Middle Miocene. Each activity is considered to have been
associated with plutonism and considerable amounts of
plutonic magmas are thought to have been emplaced into
the crust.
TECTONIC DEVELOPMENT OF THE NORTHEAST JAPAN ARC
PRE-JAPAN SEA PERIOD
Glaucophanitic metamorphic rocks are exposed in restrict-
ed localities along the western margin of the Abukuma and
Southern Kitakami terranes. This suggests the occurrence
of high pressure and low temperature metamorphism in the
Late Jurassic to Early Cretaceous (Hashimoto 1978). It is
suspected that the high-pressure metamorphism was caused
by an accretion of a subduction complex and the subse-
quent collision of oceanic plateaus in the extreme east of
the Asian continent (Minoura 1990).
Stratigraphic data imply that the terrane collision took
- Conglomerates
-v -- Calc-alkalic volcanism - Unconformity
M T s L- Sedimentary rocks - Sedimentation
10 8 6 4 2 0
Thickness (km) Fig. 8 Simplified sedimentary history of
the Southern Kitakami terrane. Alphabetical
symbols in circles show rock species of
conglomerates. The main species are indi-
cated by larger letters. G: Granitic rocks;
T: Tuff; P: Porphiritic rocks; A: Andesitic
rocks; B: Basaltic rocks; L: Limestones; C:
Cherts; M: Mudstones; S: Sandstones; H:
Holnfels. m: Mudstone beds; s: Sandstone
beds; t: Volcaniclastic beds; I: Limestone
beds. AMS: Active margin sequence; PMS:
Passive margin sequence.
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Structure and origin of the northeast Japan arc 9
c 0 15 - 13 ~a
Fig. 9 Stages in the palaeogeographic development of the northeast Japan arc during the Neogene and Quaternary periods. A: Sea area (A,: sublittoral; Ap: upper
bathyal; A,: lower bathyal); 6: Land area, C: Submarine topography (Cl: furrow; Cp: swell), D: Volcanic activity (D,: rhyolite; DE: andesite; D,: basalt; D4: basaltic pillow
lava; D,: caldera); E: Late Cenozoic stress data (El: extensional stress; E,: compressional stress).
place in Valanginian time (Minoura 1983). Every alloch-
thonous terrane experienced extensive and intensive vol-
canism in the Late Valanginian (Taketani 1987). Thus the
juxtaposition of the terranes is concluded to have been
established in the Early Valanginian.
Calc-alkalic magmas generated at shallow mantle depth
came up to the island arc crust under compressive tectonic
control. They settled as plutonic intrusions at a certain
depth of the upper crust, in Barremian to Aptian time. The
east-west increase in Sr isotope initial ratios of the plutonic
rocks (Shibata & Ishihara 1979), together with the distri-
bution of magnetite-bearing or ilumenite-series granitoids
(Ishihara 1977), presumably reflects magma generation on
the westward dipping oceanic plate (Fig. 11). A polarity in
the increasing trend of the ratios implies that the northeast
Japan arc was different in the direction of plate subduction
from the other geotectonic sections of the Japanese Islands.
Figure 12 is a summary of the pre-Tertiary tectonic
framework of northeast Japan. Each terrane is distinguish-
able from an adjoining terrane by discontinuities of
stratigraphy and structure, which suggests that a strike-slip
tectonic regime (Taira et aE. 1983; Taira & Tashiro 1987)
created the pre-Tertiary basement of the northeast Japan
arc characterized by a zonal arrangement of different types
of terranes. The exotic terranes are thought to have been
displaced along the strike-slip boundaries in the Early
Cretaceous (mainly Valanginian). Figure 13 is a simplified stratigraphy of the northeast
Japan arc and the terranes of central Hokkaido. An
ophiolitic body of the Sorachi-Yezo column (Kiminami et
al. 1986) is interpreted as oceanic fragments that have been
incorporated into the continental margin at consuming
plate boundaries. A collision of the Southern Kitakami and
Abukuma terranes with a continental plate caused the
subduction zone to jump to the oceanic side of the plateaus
(Minoura 1990). The oceanward retreat of the trench
resulted in the growth of an extensional stress regime
around abundant consuming boundaries, leading to active
volcanism with Kuroko-type ore deposition (Yamaoka
1983) in the Hauterivian (Minoura 1985).
Shortly after the terrane collision and the ensuing ig-
neous event, a major sedimentary place moved toward the
fore-arc region where the basin was floored with oceanic
fragments. Lower Cretaceous turbiditic sediments conform-
ably overlying the oceanic fragments were unconformably
overlain by Upper Cretaceous to Palaeogene thick deposits
that filled the subsiding fore-arc basin (column IV, Fig.
13). It is especially interesting that during the Cretaceous
and Palaeogene, boundaries separating every terrane were
mobile in response to the strike-slip motion between the
oceanic and east Asian plates (Minoura & Yamauchi
1989). Compressive-stress processes induced by the strike-
slip displacement of the faulted blocks formed wedge-
shaped troughs on a curved part of the faults, which were
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10 K. Minoura and A. Hasegawa
Fig. 10 Distribution of Cenozoic volcanic rocks (8) and principal strike-slip
tectonic lines (0; A-F). A: Tsushima Tectonic Line; B: Itoigawa-Shizuoka
Tectonic Line; C: Kashiwazaki-Choshi Tectonic Line; D: Tanakura Tectonic Line;
E: Karnuikotan Tectonic Line: F: Central Sakhaline-Hidaka Tectonic Line.
filled with fluvial to shallow marine clastic deposits (Fig. 13).
The fore-arc-basin sequence of the Sorachi-Yezo belt is
divisible into four stratigraphic units, every unit being
unconformable to the other (Kiminami et al. 1986). The
sedimentary processes seem to have been under the influ-
ence of plate motions, because tectonic events that oc-
curred in the northern Honshu Arc are more or less
contemporaneous with changes in the moving direction of
the oceanic plates (column VI, Fig. 13).
Intensive calc-alkalic magmatism and the simultaneous
plutonism took place during the Late Palaeocene and Early
Miocene, between 60 and 53 Ma (Shibata & Ishihara,
1979). Felsic igneous rocks having similar isotopic dates
can be found in an extensive area of east Asia (Takahashi
1983). It seems difficult to relate the magmatism simply to
a descending oceanic slab because of its large scale. Thus
another explanation for the igneous activity is required.
The separation of the North American plate from the
Eurasian plate started in the Late Cretaceous, and the
consequent spreading of the Atlantic Ocean took place at a
speed of 8-l0cm/year (Pitman & Talwani 1972). The
spreading rate was reduced to less than 4 cm/year at the end
of the Palaeocene (Solomon et aE. 1977), which resulted in
the convergence of the North American and Eurasian
plates on the opposite side of the spreading. It can be
supposed, therefore, that the magmatic event of the Late
Cretaceous to Early Palaeogene resulted from a shortening Fig. 11 Lateral variation of Sr/Sr initial ratios of Cretaceous plutonic rocks in
northeast Japan. Distribution of magnetite-bearing or ilmenite-series granitic
rocks is also shown. A: Sr/Sr initial ratios (solid arrows indicate increasing
directions); B: Distribution of magnetite-series granitic rocks; C: Transitional
zone: D: Distribution of ilmenite-series granitic rocks. Compiled from Shibata and
lshihara (1979) and lshihara (1977).
of the crust along the consuming plate boundaries, and a
collision of the continental plates.
JAPAN SEA PERIOD
The rearrangement of the northeast Japan arc and adjoin-
ing geotectonic sections in the present position seems to
have taken place in the Cenozoic, presuhably the Early
Neogene.
Maruyama and Seno (1986) proposed that the Kula
Ridge between the Kula and the Pacific plates subducted
under the Japanese Island arc in Palaeocene time. Molyb-
denum mineralization resulting from the Early Palaeogene
magmatism (Shibata & Ishihara 1979), superimposed on
the earlier polymetallic belt developed by the Late Creta-
ceous magmatism (Ishihara 1978), was probably induced by
the subduction of the ridge system (Sillitoe 1977).
Two maps in Fig. 14 show the relationships of plate
motions to plate boundaries in the Late Cretaceous and
Oligocene (modified from Pitman & Talwani [1972]). The
orogenic belts occupying most of the Cherskogo and Verkh-
oyansuk mountain ranges in east Siberia (Parfenov et al.
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Structure and origin of the northeast Japan arc 11
Fig. 12 Pre-Neogene geotectonic zoning of northeast Japan. Tectonic division
of eastern Hokkaido is provided by Kiminami et a/. (1986). Each segment is
arranged by major strike-slip displacement. ZI: Jurassic subduction complex
including mafic volcanic rocks as exotic fragments; Zll: Jurassic subduction
complex including allochthonous andesitic volcanic rocks.
1978) can be attributed to crustal shortening due to a
collision between the Eurasian and the North American
plates.
It is widely accepted that the opening of the Japan Sea
took place in the Middle Miocene, in association with the
counter clockwise rotation of the northern Honshu Arc and
the clockwise rotation of the southern Honshu Arc
(Otofuji et al. 1985). Because there was little overlap
among the fragmented crusts before or after the opening,
the rotation and simultaneous rafting of blocked land
masses are considered to have occurred along the principal
faults shown in Fig. 10. Figure 15 is the palaeogeographic
reconstruction of the Japanese Islands based on the closing
of the Japan and Yamato Basins. It has been determined
from the re-establishment of the lateral offset of metallo-
genic belts in Japan (Ishihara 1978) and Korea (Sillitoe
1977), and from palaeomagnetic data obtained by Otofuji
& Matsuda (1983) and Otofuji et al. (1985). The sedimen-
tary terrane of the Sorachi-Yezo Belt is reconstructed in parallel to northern Honshu in consideration of the offset
continuation of the Cretaceous to Palaeogene intra-arc
(Northern Kitakami) and fore-arc (Sorachi-Yezo) basin
sequences (Minoura 1989).
The oblique collision of the North American plate with
the northeast Japan block in the Eocene to Oligocene
(Minoura 1990) caused the succeeding overthrusting and
displacement of the fore-arc region of the former on to that
of the latter, resulting in the development of the Hidaka
high-P/T metamorphic body in the Late Oligocene to Early
Miocene (Arita et al. 1986). The eastward descending low-
velocity zone, recognized under the Hidaka terrane
(Miyamachi & Moriya 1984), is a possible candidate for the
latent fore-arc part of the northeast Japan arc. The oblique
convergence was followed smoothly by fragmentation of
the island arc crust and the succeeding block displacement,
which implies that the back-arc spreading was induced
under strike-slip tectonic control during the Early to Mid-
dle Miocene.
The Hawaiian seamount chain abruptly shifted the direc-
tion of growth three times in the Late Cenozoic (26, 15 and
1 Ma) to due north, deviating from the average trend (Seno
& Petersen 1984). Magmatism was activated on the east
(Andes) and west (Japan) sides of the Pacific Ocean around
these periods (Minoura 1989). Thus sharp changes in
directions of plate subduction seem to have been associated
with a great disturbance of secondary convection in the
mantle wedge and voluminous magmas were generated as a
consequence. The crust heated by ascending magmas is
expected to have been susceptible to motion in response to
strike-slip stress. The oblique collision created strike-slip
faults extending north and south on the continental plate,
through which volcanic materials originated from the mag-
mas were brought to the surface. In Middle Miocene time
(15-13 Ma) the fragmented land masses were displaced
along strike-slip faults, resulting from an increase in mobil-
ity of the lower crust heated by ascending magmas. The
block displacement and the consequent growth of a back-
arc basin can be attributed to the world-wide changes in
plate motions.
CRUSTAL STRUCTURE OF THE NORTHERN HONSHU ARC
Figure 16 presents a possible interpretation of the crustal
profile across the northeast Japan arc (see line a-A-B-b
on Fig. 2). This crustal section has been constructed using
the results from geophysical studies and adjusting the strata
so as to explain geotectonic evolution. Following Segawa
and Furuta (1978) and Yoshii (1979), lower crustal geom-
etry has been interpreted.
P-wave seismic velocity suggests that the upper crust is
principally made up of granitic intrusions and metamorphic
rocks originating from a Jurassic subduction complex,
whereas amphibolite and gabbroic rocks are main con-
stituents of the lower crust. The upper crust is subdivisible
into three lithologic units on the basis of rock types of
plutonic intrusions: Unit 1 is characterized by five kinds of
plutonic rocks resulting from the Early Cretaceous, Early
Palaeogene, Late Oligocene to Early Miocene, Middle
Miocene and Quaternary magmatism; Unit 2 includes
Lower Cretaceous and Lower Palaeogene plutonic rocks;
and Unit 3 is attended only by Lower Cretaceous plutonic
rocks.
The profile illustrates the upper crustal structure in
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12 K. Minouru und A. Hasegawa
io-
50 -
ioo-
150-
Ma
Late Cretaceous Fig. 13 Simplified stratigraphy of northern
Honshu (columns I to Ill after Minoura 1989)
and Hokkaido (columns IV and V, after Kimina-
mi et a/. 1986). I: Abukuma belt; II: Southern
Kitakami belt: 111: Northern Kitakami belt;
IV: Sorachi-Yezo belt; V: Hidaka belt. Column
VI shows relative motion between continental
and oceanic plates estimated from the calcu-
lated data of moving directions and rates of
oceanic crusts by Maruyama and Sen0 (1 986).
1 : Strike-slip movement; 2: Oblique subduc-
tion; 3: Acute-angled subduction: a: Passive
margin deposits: b: Jurassic accretionary com-
plex including mafic (ZI) and andesitic (Zll)
volcanic rocks: c: Ophiolitic complex; d: Fore-
arc basin deposits; e: Strike-slip basin depos-
its; f: Mafic to felsic volcanic rocks; g: Plutonic
rocks; h: High pressure-low temperature meta-
morphism; i: Unconformity; j: Collision of Oceanic
plateaus.
Fig. 14 Palaeotectonic maps of Eurasian and North American plates. Euler pole of plate motion, located to the northeast of Greenland in the Late Cretaceous, jumped
to east Asia in the Late Oligocene due to a collision between the Eurasian and North American plates. EP: Eurasian plate; NP: North American plate; KP: Kula plate;
PP: Pacific plate: PsP: Philippine Sea plate. Solid arrows indicate the direction of plate motion. Modified from Pitman and Talwani (1972).
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Structure and origin of the northeast Japan arc 13
Fig. 15 Schematic palaeogeographic-palaeo-
tectonic maps for the Japanese Islands. Thick
arrows in the Early Miocene map show moving
directions of blocked land-masses. See text for
discussion.
f a
I B b
I I
I I I I I I I I
krn do 300 0
TARO FAULT .
SPINEL
GARNET
PER1 DOT IT E Fig. 16 Interpretative schematic profile of the A
crust and upper mantle structure across line
a-b on Fig 2 Shown are subdivisions of the
upper crust, Unit 1 to Unit 3 Closed triangles
denote the location of Quaternary volcanoes
A Upper Cenozoic sedimentary cover, B Unit
3, C Unit 2, D Unit 1, E Gabbroic rocks with
a sheeted dyke complex, F Oceanic materials
J T Japan trench Vertical scale is exaggerated E
F
detail, showing an asymmetric nature of rift structures.
Tilted fault blocks of Cretaceous and older strata were
covered with Neogene sediments which were weakly de-
formed by east-west compressional tectonics that occurred
in the Early Quaternary (Minoura & Kosuga 1989). The
deepest part of the Jurassic subduction complex is presum-
ably composed of tectonic slices of oceanic plate facies
which were offscraped and incorporated through the devel-
opment of thrust duplexes on the subducting oceanic plate.
Because magnetic and gravity high anomalies suggest the
presence of a latent dyke-like mafic rock body beneath the
Sanriku coast, the buried mafic belt seems to reflect
oceanic materials brought into the shallow layer of the
upper crust through the Taro Fault which has been mobile
in response to oceanic plate motions since the Coniasian
(Minoura 61 Yamauchi 1989). Gabbroic rocks with a sheeted
dyke complex under the Japan Sea coast were probably
emplaced into the island arc crust in the Early Neogene due
to the opening of the Japan Sea. 0
10
20
30
40
sokm
Oceanic fragments incorporated into the continental
margin in Valanginian time were displaced through deep
fractures resulting from the oblique convergence of the
Eurasian and North American plates in the Cenozoic. This
caused an exposure of ophiolitic sequences throughout the
Sorachi-Yezo belt.
The origin and tectonic significance of the northeast
Japan arc are still speculative. A further detailed knowledge
of the crustal structure is indispensable for an understand-
ing of the geotectonic processes of plate convergence.
Super-deep drilling into the lower crust may provide the
best data to explain the origin of the arc.
ACKNOWLEDGEMENTS
The authors thank Professor A. Taira for inviting us to
express our views on the geotectonic evolution of the
northern Honshu arc. The authors also thank Drs K. Tamaki,
M. Takahashi and K. Nakamura for useful comments.
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14 K. Minoura and A. Hasegawa
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Minoura (1992) Crustal structure and origin of the northeast Japan arc.txt
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GENERATION OF CENOZOIC GRANITOIDS IN HOKKAIDO (JAPAN):
CONSTRAINTS FROM ZIRCON GEOCHRONOLOGY, Sr-Nd-Hf
ISOTOPIC AND GEOCHEMICAL ANALYSES, AND IMPLICATIONS FOR
CRUSTAL GROWTH
BOR-MING JAHN*,†, MASAKO USUKI*, TADASHI USUKI**,
and SUN-LIN CHUNG*
ABSTRACT. The island of Hokkaido is a young accretionary terrane, basically built
with a Jurassic accretionary complex and Cretaceous arc in the west ( /H11549NE Japan arc
terrane), a Cretaceous-Paleogene forearc basin and accretionary complex with theHidaka metamorphic belt in the center, and a Cenozoic island arc with Cretaceousbasement in the east ( /H11549Chishima or Kuril arc terrane). Though volumetrically small,
Paleogene and Neogene granitoids are widespread in central Hokkaido (Hidaka Belt).Granitoids are the most representative component of the continental crust, so in thiswork we aimed to study the mode of generation and source characteristics of thesegranitoids in order to assess the crustal composition of Hokkaido and examine thegeneral problem of continental growth. New zircon geochronology on nine graniticand one gabbroic rocks from the Hidaka Belt reveals three distinct magmatic episodes,two in the Eocene at 45–46 Ma (3 granites), and 37.0 /H115500.5 Ma (1 granite), and one in
the Miocene at 18 to 19 Ma (5 granites and 1 gabbro). The Miocene episode representsthe most important granitic emplacement in Hokkaido. The early Eocene zircon agesof 45 to 46 Ma are identified for the first time for granitoids that occur in the northernpart of the Hidaka Belt. The zircon age of 37 Ma for a granite from Shirataki is ratherrare in Hokkaido, but similar ages had been reported for a tonalite and a granite fromthe Hidaka metamorphic belt. Geochemically, all granites are slightly peraluminousbut not S-type, and they possess volcanic arc granitoid characteristics. Their REEdistribution patterns are typically “granitic,” showing fractionated patterns with LREEenrichment and distinct negative Eu anomaly. The whole-rock isotopic signatures[I
Sr/H115490.7044 to 0.7061; /H9255Nd(t)/H11549/H11545 1.0 to /H115454.7; TDM-1/H11549400-1000 Ma] reveal their
largely juvenile characteristics. This is corroborated by the zircon Hf isotopic composi-tions [ /H9255
Hf(t)/H11549/H11545 8t o/H1154519]. The Eocene granites were most probably generated by
melting of subducted accretionary complex in a prolonged period from 46 to 37 Ma insupra-subduction zone; whereas the Miocene granites were also generated by meltingof accretionary complex in a back-arc rifting setting. In both cases, the involvedaccretionary complex was probably dominated by the mantle-derived lithologicalcomponent with little Paleozoic or older crustal material. Hokkaido provides anexcellent example of juvenile crust addition to the continental crust.
Key words: Accretionary orogen, accretionary complex, Japanese Islands, Hok-
kaido, zircon dating, Sr-Nd-Hf isotope tracers, granitoids, Nipponides, crustal growth,juvenile/recycled crust, Central Asian Orogenic Belt (CAOB)
introduction
The Japanese Islands represent a Phanerozoic subduction-related orogen devel-
oped along the western Pacific convergent margin. The formation of the JapaneseIslands has been taken as the classic model for accretionary orogeny (for example,Cawood and others, 2009). According to Maruyama and associates (Maruyama, 1997;Maruyama and others, 1997), the most important cause of the orogeny is thesubduction of an oceanic ridge, by which the continental mass increases through thetransfer of granitic melt from the subducting oceanic crust to the orogenic belt.
* Department of Geosciences, National Taiwan University, Taipei 106, Taiwan
** Institute of Earth Sciences, Academia Sinica, Nangang, Taipei 115, Taiwan
†Corresponding author: (bmjahn@ntu.edu.tw)[American Journal of Science, Vol. 314, February, 2014 ,P .704–750, DOI 10.2475/02.2014.09]
704
Sengor and Natal’in (1996) named the orogenic complex the “Nipponides,” and
pointed out the resemblance in orogenic style between Japan and the Central AsianOrogenic Belt (CAOB). Using the newly acquired and literature Sr-Nd isotopic data,Jahn (2010) tested the models made by the above authors. The test reveals that a largeproportion of the granitoids from SW Japan in fact show high initial
87Sr/86Sr ratios,
negative εNd(T) values and Proterozoic Sm-Nd model ages. These data are in strong
contrast with those of two celebrated accretionary orogens, the CAOB (for example,Jahn, 2004) and Arabian-Nubian Shield (ANS; for example, Stern, 1994; Eyal andothers, 2014), but are quite comparable with those observed in SE China and Taiwan(Jahn and others, 1990; Chen and Jahn, 1998; Jahn, 2010). This raises questions aboutthe bulk composition of the continental crust in SW Japan, or the type of materialaccreted in accretionary complexes. The finding of Jahn (2010) also negates thehypothesis that the Nipponides contains very few fragments of older continental crust.It appears that the subduction-accretion complexes in SW Japan represent only theupper portion of the bulk crust, which is probably underlain by a Proterozoicbasement.
A continuous study on other parts of the Japanese Islands reveals that the crustal
development in NE Japan ( /H11005NE Honshu and Hokkaido) is quite distinguished from
that of SW Japan (fig. 1). In NE Japan the granitoids show the lithological types (moreTTG, or tonalite-trondhjemite-granodiorite suite, and adakitic rocks) and Sr-Ndisotopic compositions with significantly more “juvenile” signatures. Since graniticrocks are commonly generated in the P-T conditions of middle to lower crust, they areideal materials to be used to probe the nature and architecture of the middle to lowercontinental crust. Besides, the radiogenic isotopic compositions of granitoids preservethe crustal history of their protoliths and constrain their formation time. In this paper,we employ the conventional geochemical and isotopic tracer techniques, together withzircon dating and Hf isotope analyses, to examine the petrogenesis of the Tertiarygranitoids from Hokkaido. We will then compare the results with that of the massivegranitoids from SW Japan and other celebrated accretionary orogens (for example,CAOB and ANS), and discuss the implications for crust growth and the tectonicevolution of the Japanese accretionary orogens.
general geologic setting of japan and hokkaido
The evolution of the Japanese Islands results from the interaction of four tectonic
plates: the Eurasian, Philippine Sea, Pacific and Okhotsk. The southwestern part ofJapan (SW Japan) is an eastern margin of the Eurasian Plate, but separated from theAsian continent by the Japan Sea. The northeastern part of Honshu and Hokkaido(/H11005NE Japan) belong to the Okhotsk plate. The Okhotsk plate was formerly consid-
ered as a part of the North American Plate, but recent studies indicate that it is anindependent mini-plate, bounded on the north by the North American Plate (Senoand others, 1996; Apel and others, 2006). The Philippine Sea plate is subductingnorthwestwards at a rate of 4 to 6 cm/a under SW Japan along the Nankai Trough andRyukyu trench. The Pacific plate is subducting at a rate of 9 to 10 cm/a beneath NEJapan, with its leading slab reaching a depth of 660 km underneath the area of Beijing,China, as revealed by a tomographic study (Zhao and others, 2007, 2011; see also areview by Isozaki and others, 2010).
An incipient subduction zone appears to be developed in the eastern Japan Sea
(Nakamura, 1983; Tamaki and Honza, 1985). Its onland extension, the Itoigawa-Shizuoka Suture (fig. 1), separates the Japanese Islands in two parts, NE and SW Japan.However, this subduction zone is not universally accepted as the geophysical evidenceis still not so convincing. On the other hand, many workers consider that the truegeological or tectonic boundary between NE and SW Japan in the Pre-Tertiary time islocated within NE Honshu, termed as “the Tanakura Tectonic Line” (fig. 1).705 B.-M. Jahn & others 705
The predominance of accretionary complexes and the association of detached
continental fragments in Japan suggest that the Japanese Islands have developedmainly through convergence between oceanic and continental plates along activemargins (Isozaki, 1997; Isozaki and others, 2010). Isozaki (1996) stated that severalmajor oceanic plates have subducted beneath the South China Block margin, leavingmore than 10 distinct accretionary complex (AC) belts (now reduced to 9 AC belts,based on the latest reappraisal of the geotectonic framework of Japan by Isozaki andothers, 2010). All the AC belts occur as thin subhorizontal fault-bounded geologicunits, that is, nappes, and show a clear downward and oceanward younging polarity(Isozaki and Itaya, 1991; Isozaki and Maruyama, 1991). Numerous oceanic fragmentsderived from subducted oceanic plates, including deep-sea sediments and seamountbasalts and reef limestone, were accreted to Japan. According to Isozaki and others(2010), five orogenic phases had occurred in the last 500 Ma, namely, at 450 MaFig. 1. Index map of Japan showing major tectonic units discussed in this paper. NE Japan is separated
from SW Japan by a major fault or tectonic boundary, which is controversially represented by theItoigawa-Shizuoka or Tanakura Tectonic Line. SW Japan comprises five tectonic belts: three granitic belts(Sanin, Sanyo and Ryoke) are separated by two accretionary belts (Sanbagawa and Shimanto) by the MedianTectonic Line. In NE Japan, granitic rocks mainly occur in the Kitakami and Abukuma Mountains.706 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
(Oeyama), 340 Ma (Renge), 240 Ma (Akiyoshi), 140 to 130 Ma and 80 to 60 Ma. The
Permo-Triassic event in Japan was thought to be related to the continental collisionbetween the North and South China Blocks as recorded in the Dabie-Sulu terrane ofChina (Oh, 2006; Isozaki and others, 2010), or due to the collision of a Proto-Japanblock with the Eurasian margin (de Jong and others, 2009). In summary, the model ofaccretionary orogeny developed for the Japanese Islands ( /H11005Miyashiro-type orogeny,
Maruyama, 1997) underlines the prime role of continuous ocean-floor and episodicocean-ridge subduction.
Geology of Hokkaido . The island of Hokkaido is a young accretionary terrane
with little or no rocks of Paleozoic ages and older. Kiminami and others (1986)presented the first tectonic framework of the island’s evolution, which was followed byother workers (for example, Komatsu and others, 1992). According to Ueda andothers (2000) and Ueda (2005), Hokkaido comprises five roughly N-S runningtectonic units or orogenic belts, from west to east (fig. 2): (1) the Oshima Belt, aJurassic accretionary complex and an overlying Cretaceous arc; (2) the Sorachi-YezoBelt, a Cretaceous-Paleogene forearc basin and accretionary complex; (3) the HidakaBelt, a Paleogene arc complex in the north and the Hidaka metamorphic belt in thesouth; originally, the Hidaka Belt was defined to contain a Hidaka Supergroup (clasticaccretionary complex) and a Hidaka Metamorphic Belt (metamorphosed accretionarycomplex, up to granulite facies); (4) the Tokoro Belt, a Cretaceous and Paleogeneaccretionary complex; and (5) the Nemuro Belt, a Cretaceous and Paleogene arc/forearc complex (fig. 2). Tectonically, the Oshima Belt has been considered as thenorthern extension of NE Honshu and formed the same tectonic collage withSikhote-Alin of the Russian Far East ( /H11005Honshu-Sikhote-Alin Tectonic Collage), and
the Sorachi-Yezo and Hidaka Belts formed a second tectonic collage or continental-margin arc with Sakhalin ( /H11005Sakhalin-Hokkaido Tectonic Collage; Rodionov and
others, 2011). This arc was interpreted as having formed during subduction of theancestral Pacific plate (Izanagi). The Tokoro and Nemuro belts belong to the Kuril ArcTerrane.
The accretionary complexes of the Oshima, Sorachi-Yezo and Hidaka Belts show a
generally eastward younging polarity formed by westward subduction (Ueda andothers, 2000; Kawamura, 2004). U-Pb SHRIMP dating of detrital zircon from theOshima sandstone revealed some Precambrian ages of 1.88 and 2.5 Ga, hencesuggesting that the Oshima sandstone had a share of clastic source from Precambrianterranes of the Asian continent or the Sino-Korean Craton (Kawamura and others,2000).
The Sorachi-Yezo Belt comprises four tectonic units (GSJ, 2010): (1) the Sorachi
Group, in the western part of the belt, composed essentially of greenstones in the lowerpart and greenstone, chert and pyroclastic rocks in the upper part; (2) the Yezo Group,to the east of the Sorachi group, characterized by a Cretaceous forearc basin sequencedominantly of marine siliciclastic deposits; (3) the Kamuikotan Zone, consisting ofCretaceous accretionary complexes, which have undergone various grades of metamor-
phism from the blueschist to epidote amphibolite facies; and (4) the Idonnappu Zone,consisting also of Cretaceous accretionary complexes, but only feebly metamorphosed.Kimura and others (1994) proposed that the greenstones of the Sorachi Grouprepresented the remnants of an accreted Late Jurassic oceanic plateau formed on theIzanagi (Paleo-Pacific) plate based on the large volume of basaltic flows and hyaloclas-tic deposits. However, it has also been proposed that the Sorachi-Yezo belt represents anormal oceanic crust (Niida and Kito, 1986) or a marginal basin crust (Takashima andothers, 2002). As a whole, Hokkaido is characterized by the latest Cretaceous to earlyPaleogene rapid growth of accretionary complex and exhumation of high pressuremetamorphic rocks in the northwestern Pacific margin (Kimura, 1994).707 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Geologic map
modified from
Yamada and others (1990)
Tectonic units after
Ueda and others (2000)
and Ueda (2005)
Fig. 2. General geologic map of Hokkaido (modified after Yamada and others, 1990). The division of
tectonic units follows Ueda and others (2000) and Ueda (2005).708 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
The Hidaka Belt in central Hokkaido is characterized by the vast distribution of
Cretaceous accretionary complexes and Tertiary metamorphic belts and ophiolites(for example, Osanai and others, 1991, 1992). The Hidaka Supergroup consists mainlyof sandstone and shale, with minor chert and volcanic tuffs. Basaltic lava is also foundto erupt on or intruded into unconsolidated mudstone (Miyashita and Katsushima,1986; Kiminami and others, 1999). Based on microfossil studies, these sediments weremostly deposited in the Paleogene. The Supergroup was intruded by granitoids ofEocene to Miocene ages. All the samples analyzed in this work came from this belt.
The Hidaka Metamorphic Belt (HMB) in south-central Hokkaido consists of
high-angle east-dipping thrust sheets composed of metamorphic (pelitic-psammiticrocks, intermediate and mafic rocks), igneous (layered gabbro, massive gabbro-dioriteand granitoids), and alpine-type ultramafic rocks (Arai and Takahashi, 1989; Arai,1994; Shimura and others, 2004). In fact, the southern part of the Hidaka belt wasconsidered to have formed by upthrusting of the eastern main block toward thewestern block due to oblique subduction of the Pacific Plate along the Kuril Trenchsince Miocene (Kimura, 1986). Thus, some workers considered that the zone along theHidaka Main Thrust is a distinct suture between the western (Eurasian plate) andeastern (Okhotsk Plate) blocks (for example, Iwasaki and others, 2004). On the otherhand, Ueda (2005) considered that the boundary between the Kurile and NE Japan arcterranes is situated to the east of the Hidaka Belt.
The base of the HMB is the Hidaka Main Thrust, where the metamorphic and
igneous rocks are mylonitized with dextral shear sense. Almost all igneous rocks areintruded into the metamorphic layers as syn-metamorphic suites. The HMB is thoughtto represent an eastward dipping island-arc type crustal section (Komatsu and others,1989, 1994; Owada and others, 2003). Based on the reconstruction of Shimura andothers (2004), the unexposed “lowermost part” of the “Hidaka crust,” from ca.23 km
to the Moho (30 km?) is probably composed of mafic granulites as inferred frompetrological studies. The “lower Hidaka crust” (15-23 km) is likely represented bygarnet-2-pyroxene mafic granulite, garnet-orthopyroxene aluminous granulite andgabbro. The mafic granulites and gabbros show MORB-like composition (Maeda andKagami, 1994, 1996; Mikoshiba, 1999).
The Hidaka Mountains show a positive gravity anomaly zone, which was probably
due to the tectonic uplift of the basement rocks that occurred during the collision ofthe Kuril forearc sliver with the northern extension of the Honshu arc (Kimura, 1986;Taira, 2001). In addition, the crustal thickness of the Hidaka Mountains is ca.30 to 50
km (Ogawa and others, 1994), which is the thickest in Hokkaido.
The celebrated Horoman Peridotite occurs in the southernmost of the Hidaka
belt. It is a fault-bounded mantle slice of 8 km /H1100310 km /H110033 km, emplaced at ca.23 Ma
(Rb–Sr isotopes on a phlogopite-bearing spinel lherzolite, Yoshikawa and others,1993). The peridotite consists of several lithological sequences of plagioclase lherzolite–lherzolite–harzburgite /H11006dunite–harzburgite–lherzolite–plagioclase lherzolite (for
example, Takahashi, 1991; Takazawa and others, 1999, 2000). As the HoromanPeridotite is not directly related to the present work, no further details will be given.Similarly, the two tectonic units of the Kuril Arc, the Tokoro and Nemuro belts, areunrelated to the petrogenesis of the granitoids concerned, their description is notfurther presented. We proceed to introduce the geological setting and essential pointsrelevant to the studied granitoids.
occurrence of granitoids
Though volumetrically small, Cenozoic granitoids are widespread in the Hidaka
Belt. They are scattered in the axial belt of central Hokkaido for an area of 300 km N-Sand 60 km E-W (Ishihara and Terashima, 1985). The granitoids commonly crop out tothe east of the high-grade metamorphic belt in the south, but scattered widely in the709 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
non-metamorphosed Hidaka Supergroup in the north. Spatially, the granitoids do not
accompany coeval volcanic rocks, but are closely associated with the contemporaneousgabbroids.
The Cenozoic granitoids of the Hidaka Belt are mostly fine-grained massive I-type
granitoids (Ishihara and Terashima, 1985). Since they are associated with gabbroids,especially in the westernmost zone, the plutonism is of a bimodal nature. S-typegranitoids also occur, and both S- and I-type granitoids contain ilmenite but notmagnetite, so they belong to the ilmenite series (Ishihara and others, 1998). Ishihara(2007) considered that the granitoids were generated within the accretionary complexof the Hidaka Supergroup and its basement, by heat provided by these gabbroids.
Tonalitic granitoids, commonly known to be generated by partial melting of mafic
sources, such as amphibolites, at lower crustal conditions, are conspicuously present inabundance in the Hidaka Metamorphic Belt. Based on the ASI (aluminum saturationindex) or A/CNK ( /H11005Al
2O3/(CaO /H11001Na2O/H11001K2O) molar ratio), many of them have
been classified as “S-type tonalities,” and the peraluminous variety (S-type) predomi-nates over the metaluminous I-type (Shimura and others, 2004). Furthermore, thetonalities, especially the S-type, entrain a large number of enclaves of various rock-types, including para-granulites (gt-opx-bi granulite and gt-opx-cord granulite), maficgranulites, and meta-harzburgite. Mafic enclaves comprise gabbro, hornblende amphi-bolite and mafic granulites (opx granulite and two-px granulite). Enclave-melt reac-tions are commonly observed.
Kemp and others (2007a) studied a suite of rocks including opx-bearing granu-
lites, tonalites and gabbros from the Hidaka metamorphic belt and attempted to clarifythe magmatic-metamorphic connection in this area. Their zircon geochronologyrevealed that the granulite and amphibolite facies metamorphism and the emplace-ment of garnet-opx tonalites and gabbros took place at about 19 Ma; whereas ahornblende-tonalite and a granite were emplaced at 37.5 /H110060.3 Ma. With these zircon
age data, they concluded that the Hidaka metamorphic belt has recorded a two-stageevolution, with the first stage of supra-subduction zone magmatism in late Eocene ( ca.
37 Ma) and the second stage of back-arc extension in the Miocene ( ca.19 Ma). The
second stage coincides with the opening of the Japan Sea, and the tectonic activityresulted in the granulite metamorphism and generation of grt-opx tonalite andgabbro, probably all related to the underplating of basic magma.
sampling localities and sample description
The granitoid samples of this study were collected from the Hidaka Belt of central
Hokkaido, in a N-S traverse along about 143°E (fig. 3). A simple petrographicdescription of all the samples is given in table 1. Note that the rock types assigned forthe granitoids were determined after the QAPF classification scheme of Streckeisenand Le Maitre (1979) based on normative abundances of quartz and feldspars. Themodal analysis indicates that in addition to quartz and feldspars, biotite is present in alland hornblende in most granitoid samples. The only gabbro sample is composed ofplagioclase, orthopyroxene and clinopyroxene.
analytical methods
Zircon U-Pb Geochronology and Lu-Hf Isotopic Analysis
Zircon grains were separated from samples of about 1 to 3 kg using the conven-
tional heavy-liquid and magnetic separation techniques at the Langfang MineralSeparation Laboratory, near Beijing. Cathodoluminescence (CL) images were taken atthe Beijing SHRIMP Center, Institute of Geology, Chinese Academy of GeologicalSciences, for examination of zircon internal structures and for selection of analyticalspots.710 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Zircon U–Pb isotopic analyses were performed using a New Wave UP213 laser
ablation system combined with an Agilent 7500s quadrupole ICPMS (inductivelycoupled plasma mass spectrometer) at the Department of Geosciences, NationalTaiwan University (NTU-Geosciences). The LA-ICPMS operating conditions andanalytical procedures were the same as those reported in Chiu and others (2009). WeFig. 3. Sampling localities of the granitoid samples from central Hokkaido. The Yamabe and Perari-
yama intrusions are of granite porphyry. The new zircon ages obtained in this study are shown next to thesample numbers. Map based on GSJ, AIST, editor (2003), Maeda and others (1986), Editorial committee ofHokkaido (1990), Editorial committee of Geology of Japan (2005), Osanai and others (2006, 2007), Suetake(1997), and Nakagawa (1992).711 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
have followed a common practice in reporting zircon ages of young, particularly
Cenozoic, rocks; for example, Wen and others (2008) and Chiu and others (2009).Generally, precise measurement of
207Pb/235U and207Pb/206Pb ratios is feasible for
Precambrian zircons, but not for very young zircons, due to the fact that in youngzircons
235U comprises less than 1 percent of natural U, thus little207Pb can be
produced in zircons (for example, Ireland and Williams, 2003). For this reason, theweighted mean of pooled
206Pb/238U ages are taken to represent the crystallization
ages of the dated samples. The206Pb/238U ages are reported with uncertainties at
two-standard deviation (2 /H9268) or 95 percent confidence level.
In-situ Lu-Hf isotopic analyses of zircon were performed using a multi-collector
ICP-MS (Neptune), also at NTU-Geosciences. A New Wave UP193FX laser ablationsystem was used for spot vaporization. The Lu-Hf isotope analyses were done on thesame zircon grains that were previously analyzed for U-Pb dating. Ablation time wasabout 26 s for each measurement with a beam diameter of ca.40/H9262m, an 8 Hz
repetition rate, and energy of 100 mJ. The detailed descriptions for the analyticaltechniques can be found in Wu and others (2006) and Xie and others (2008). TheHarvard reference zircon 91500 and Australian Mud Tank carbonatite zircon wereused as secondary standards for data quality assessment. During the data acquisition ofthis study,
176Hf/177Hf ratio of 0.282511 /H1100625 (2 /H9268,n/H1100539) for Mud Tank and
0.282293 /H1100622 (2 /H9268,n/H1100516) for 91500 were obtained. These values are in good
agreement with those obtained by solution and ICPMS methods reported in theliterature (Goolaerts and others, 2004; Woodhead and others, 2004; Woodhead andHergt, 2005; Griffin and others, 2006; Wu and others, 2006).
Major and Trace Element Analyses
All major and trace element analyses were performed at NTU-Geosciences.
Samples were crushed in a stainless steel jaw crusher and then powdered in an agateTable 1
Petrographic description of granitoid and gabbro samples from Hokkaido, Japan712 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
mill. Major elements were determined by X-ray fluorescence (XRF) spectroscopy on
fused glass beads, using a Rigaku RIX-2000 spectrometer. For trace-element analyses,about 200 mg of powdered sample was dissolved in a mixture of HF and HNO
3(2:1) in
a screw-top Teflon beaker (Savillex) for 5 to 7 days at /H11011100 °C. This was followed by
evaporation to dryness, refluxing in 6N HCl and drying twice, and finally re-dissolutionin 1N HCl. The procedure was repeated until complete dissolution. The final solutionwas split in two parts; a small aliquot (about 10%) was used for subsequent traceelement analysis by ICP-MS, and the rest for further chemical separation of Sr and Ndfor isotopic analysis using a thermo-ionization mass spectrometer (TIMS). Traceelement analysis was performed using an Agilent 7500s. The standard referencematerials used for trace element analyses are AGV-2, BCR-2, BHVO-2, BIR-1 andDNC-1. The details of analytical procedures may be found in Lin and others (2012).Analytical errors are 0.5 to 3 percent for major elements and 1 to 10 percent for traceelements, depending on the concentrations.
Whole-Rock Sr-Nd Isotopic Analyses
For Sr-Nd isotopic analysis, the chemical preparation and mass analysis were
performed at Institute of Earth Sciences (IES), Academia Sinica. Approximately 150 to175 mg of rock powder was dissolved using a HF–HNO
3(2:1) mixture in a screw-top
Teflon beaker for 5 to 7 days at /H11011100 °C. This same procedure was followed by
evaporation to dryness, refluxing in 6N HCl and drying twice, and then dissolution in1N HCl. The procedure was repeated until complete dissolution. Chemical separationwas carried out using the conventional ion exchange techniques. Strontium and REEswere separated in polyethylene columns with a 2.5 ml resin bed of AG50W-X8, 100 to200 mesh. Strontium was further purified through 1 ml resin bed of AG50W-X8, 100 to200 mesh. Neodymium was separated from other REEs on 1 ml polyethylene columnsusing Eichrom Ln resin (Ln-B25-A) as a cation exchange medium. Sr and Nd isotoperatios were measured using a Finnigan MAT 262 and a TRITON mass spectrometer.For the isotopic measurement, Sr was loaded on a single Ta filament with H
3PO4and
TaF5; but Nd was loaded on a Re filament with H3PO4and measured using a
double-Re-filament configuration. The effect of mass fractionation in Sr and Ndisotopic measurements was corrected by normalizing to
86Sr/88Sr/H110050.1194 and
146Nd/144Nd/H110050.7219, respectively. Analyses of NBS 987 Sr and JMC Nd standard
throughout the period of analysis yielded86Sr/87Sr/H110050.710238 /H110060.000016 (2 /H9268) and
143Nd/144Nd/H110050.511812 /H110060.000007 (2 /H9268). Procedural blanks were ca.330 pg Sr and
300 pg Nd. Within-run precision, expressed as 2 /H9268m, was better than 0.000010 for both
Sr and Nd. The procedures of chemical separation and mass analysis can be found inJahn and others (2009).
analytical results
Zircon U-Pb Data
The CL images of the ten analyzed zircon samples are shown in figure 4. The sizes
of zircon grains could be estimated from the round laser-abraded spots that have adiameter of about 50 /H9262m. Zircon grains are in general prismatic and euhedral. A total
of 104 images were taken. Since all the images are very similar, only two images fromeach sample are displayed in this figure. All zircon crystals show simple internalstructure with clear oscillatory zonings, thus their magmatic origin can be certified.The results of U-Pb isotopic analyses are given in table 2. The errors for individual spotanalyses are quoted at 1 /H9268, whereas those for the weighted mean ages represent 2 /H9268
(95% confidence level).
Figure 5 shows a plot of Th/U ratios vs U concentrations in zircon samples. The
variation in both parameters are quite impressive; the U concentrations vary from less713 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
than 100 ppm to 2000 ppm or more, and most Th/U ratios are greater than 0.5. Since
metamorphic zircons commonly have low to very low Th/U ratios ( /H110210.1; Hoskin and
Black, 2000), the present Th and U concentration data corroborate the magmaticorigin of the zircon crystals of the Hokkaido granites.
Figure 6 illustrates the U-Pb isotopic compositions in the Concordia diagrams. We
underline that almost all individual analyses fall on or near the Concordia. For theobvious reason that the ratios of
206Pb/238U are more precisely measured than those of
207Pb/235U and207Pb/206Pb (table 2), the weighted mean values of206Pb/238U dates
are taken to be the crystallization ages of the analyzed zircon crystals. The obtainedages fall in three groups: (1) 45 to 46 Ma (3 granite samples), (2) 37.0 /H110060.5 Ma (1
granite sample), and (3) 18 to 19 Ma (5 granite samples and 1 gabbro sample).
Whole-Rock Geochemical Data
The chemical analyses of the granitoids are presented in table 3. The principal
characteristics are illustrated in figures 7 and 8. In the Q’-ANOR classification schemeof Streckeisen and Le Maitre (1979), all the Hokkaido granitoids fall in the class of“granite” (monzogranite /H11001syenogranite) with only one exception, sample SRK-2, in
granodiorite (fig. 7A). The granitoid samples show a range of SiO
2contents from ca.
65 to 75 percent, and K2O from 2.7 to 5.0 percent (table 3). In the A/NK vs A/CNK
diagram (fig. 7B), most samples are shown slightly peraluminous; all but one sample(AB-1) have A/CNK ratios less than 1.1, which is the value suggested by Chappell andWhite (1992) to be the boundary between I- and S-type granitoids. Consequently, mostgranitoids may be considered as I-type granitoids, but not S-type as commonly referredto in the literature.
Fig. 4. Representative CL images of the dated granitoids and a gabbro. The round spot size is about 50
/H9262m (diameter).714 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 2
Zircon U-Pb isotopic compositions, Th-U concentrations and calculated ages715 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 2
(continued)716 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 2
(continued)717 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 2
(continued)718 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 2
(continued)719 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 2
(continued)720 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 2
(continued)
c/H11005core; r /H11005rim721 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Chondrite-normalized REE patterns are shown in figures 8A and 8B. They are
typical of granitic rocks, with light REE enrichment and conspicuous negative Euanomaly. The REE abundances of the older Eocene granitoids (fig. 8A) may be slightlylower and less fractionated than those of the younger Miocene granitoids (fig. 8B), butthe difference is quite subtle. The only gabbro sample exhibits a quasi-flat REE patternwith about 10x chondritic abundances. In a sense it looks like an atypical N-MORB withsmall LREE depletion.
Primitive-mantle-normalized spidergrams of the granitoids are shown in figures
8C and 8D. In the spidergrams the trace elements are arranged in the ascending order,from left to right, of their compatibilities with the basaltic liquid. Nevertheless, theapplication of such diagrams to granitic rocks also serves to identify fractionation ofparticular mineral phases during the generation and differentiation of granitic liquids.Figure 8C shows that in the Eocene granitoids, depletion or negative anomaly isobserved in Nb-Ta, Sr, P, Zr and Ti. The phenomenon of the “TNT (Ti-Nb-Ta)anomaly” is most characteristic of granitic rocks, island arc volcanics and the continen-tal crust in general.
The spidergrams of the Miocene granitoids are shown in figure 8D. The general
enrichment/depletion patterns are grossly similar to those of the Eocene granitoids.The only gabbro shows negative anomalies in Nb and Ta, but positive anomaly in Sr.Such an elemental distribution may favor an interpretation of its origin in an island arcsetting, but not in mid-ocean ridge.
Whole-Rock Sr-Nd Isotopic Data
The whole-rock Sr and Nd isotopic analyses are given in table 4, and further
illustrated in figure 9A. The Rb concentrations in all samples range from 50 to 143ppm, which are “normal” for granitic rocks. However, the Sr contents vary from ca.50
to 190 ppm, which are somewhat lower than the normal range of granitic rocks. TheUK-1ICH-1
UTT-2
AB-1OTC-1
OTC-2NS-1
SAH-1SRK-1ST-1
0.00.51.01.52.02.53.03.54.0
Th
U
0 500 1000 1500 2000 2500
U (ppm)Magmatic ratios
Metamorphic ratios
Fig. 5. Th and U concentrations in zircon crystals of the Hokkaido granitoids. The generally high Th/U
ratios ( /H110220.5) suggest that the zircon crystals are of magmatic origin.722 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
calculated initial87Sr/86Sr ratios (ISrvalues) range from 0.7044 to 0.7061; and the
age-corrected initial143Nd/144Nd ratios, expressed as εNd(t) values, are all positive,
ranging from /H110011.0 to /H110014.7. Single-stage Sm-Nd model ages are between 400 to 1000
Ma (table 4).
Fig. 6. U-Pb Concordia diagrams for the zircon grains from the Hokkaido granitoids and a gabbro.723 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
In figure 9B, the literature Sr-Nd isotopic data of Cenozoic felsic-intermediate
magmatic rocks are also shown for comparison. The isotopic data of Miocene rhyoliticrocks are predominant and they refer to those occurring in four areas—northernHokkaido (4-15 Ma; Takagi and others, 1999), central Hokkaido (15-17 Ma; Furukata
Fig. 6. (continued).724 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
and others, 2010), SW and NE Hokkaido (2-18 Ma; Takanashi and others, 2011, 2012).
Note that all the rocks from the Hidaka metamorphic belt show positive εNd(t) values
except for four tonalitic rocks, two of them defined as tonalitic xenoliths and two asS-type tonalities (shown in blue diamonds in fig. 9B; Owada and others, 2006). The
Fig. 6. (continued).725 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
entire data set also shows an anti-correlation between εNd(t) and ISrvalues. Note that
the ensemble of our new granitoid data points (shown in red solid dots, fig. 9B)appears to lie slightly above the bulk array.
Fig. 6. (continued).726 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Zircon Hf Isotopic Data
The zircon Hf isotopic compositions of the Hokkaido granitoids are given in table
5, and further illustrated in figures 10A and 10B. Note that all individual samples havea range of ε
Hf(t) values; for example, sample UK-1 ( ca.45 Ma) shows a range of εHf(t)
Fig. 6. (continued).727 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
from /H110018.9 to /H1100117.3, and ST-1 (37 Ma) from /H1100111.0 to /H1100118.1 (table 5). The ranges are
beyond the analytical uncertainty of 0.5 to 1.0 epsilon unit. In a study of zirconchemistry and Hf isotopic compositions from two igneous complexes (Pingtan andTonglu) in SE China, Griffin and others (2002) observed that a large variation in
176Hf/177Hf (up to 15 εHfunits) was found between zircon grains of different growth
stages within a single rock, and between zones within single zircon grains (up to 9 εHf
units). Such variation suggests that each of the observed magmas in both complexes
developed through hybridization of two or more magma batches with differentsources. They conclude that this mixing has produced similar Sr and Nd isotopiccompositions in the different rock types of each complex, but the zircons havefunctioned as “tape recorders” and preserved details of the assembly of the differentmagmas. We agree with the above observation and interpretation of Griffin and others(2002), however, we like to offer a supplementary explanation as follows.
Zircon crystallizes at a given time in an evolving granitic magma that would likely
preserve the chemical and isotopic compositions in equilibrium with the magma atthat time. The bulk composition of a magma chamber would change throughfractional crystallization, but such closed system chemical fractionation will probablynot modify the isotopic compositions. A change of isotopic compositions could only beachieved through an open-system behavior, such as influx of a foreign magma orassimilation of country rocks in the magma chamber. Zircon crystals formed at anygiven stage would register the Hf isotopic composition of the evolving magma at thatstage. Since zircon crystals do not grow at the same time, the individual grains from asingle rock could have recorded the Hf isotopic compositions of different stages ofmagma evolution.
Kemp and others (2007b) conducted an elaborate study of Hf isotopic change
with magma generation and evolution of the I-type granites from the Lachlan Fold Beltof Australia. They analyzed U-Pb, Hf and O isotopic compositions in zircons to revealthe nature of the crustal component. They reached a novel conclusion that the I-typegranites were in fact formed by the reworking of sedimentary materials by mantle-likemagmas, but not by remelting of older metamorphosed igneous rocks as widelybelieved. Nonetheless, the authors also concluded that I-type magmatism criticallyinvolves continental growth, this being camouflaged to some extent by the non–mantle-like isotope ratios of the bulk rocks. The overall proportion of juvenile material addedby the Lachlan I-type suites was between 85 percent and 50 percent in differentplutons.
Despite the large range of ε
Hf(t) values observed in the Hokkaido granitoids (fig.
10A), the entire data set shows that all of them is exclusively positive, similar to thewhole-rock ε
Nd(t) values. The two-stage Lu-Hf model ages are shown in figure 10B. All
model ages are younger than 550 Ma, and a few give negative or future age values.Similar to the whole-rock Sm-Nd model ages, the zircon Hf model ages are consistentwith their juvenile characteristics. No Precambrian heritage is identified.
discussion
Significance of the New Zircon U-Pb Ages and Literature Age Data
Our new zircon age data and those from Kemp and others (2007a) indicate three
distinct intrusive events in Hokkaido at ca.45, 37 and 18 Ma. In order to reach a better
understanding of the significance of the newly obtained zircon ages and the tectonomag-matic evolution of Hokkaido, we compiled the available age data of Tertiary plutonicrocks from the literature, and they are summarized in a histogram (fig. 11). Among the108 dates reported in the literature, a half of them (54) were obtained on biotite by theconventional K-Ar method, and a quarter (29) were fission-track (FT) dates on zircon(18) and apatite (11). Prior to the present work, zircon U-Pb dates are rare (total /H110056),728 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 3
Chemical compositions of granitoids and a gabbro from Hokkaido729 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 3
(continued)730 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
and Kemp and others (2007a) gave 5 out of 6 dates. As demonstrated earlier, the clear
magmatic zoning, the range of Th/U ratios ( /H110220.5) in zircon crystals and the simple
clustering of concordant or near-concordant data points provide strong evidence for
the zircon crystallization in magmatic liquids. The three distinct age groups mustrepresent three significant granitic intrusive episodes at 45, 37 and 18 Ma (fig. 11).
The majority of the K-Ar and fission-track (FT) dates can in principle be inter-
preted as the time of magmatic cooling but not the time of magmatic emplacement.This is probably true for the cases of Otchube (U-Pb age /H1100518, K-Ar (biotite/WR) /H11005
16.5/16.0 Ma), Aibetsu (U-Pb /H1100545.5, FT (zircon) /H1100538.9, FT (apatite) /H1100516.5 Ma), and
Nissho-toge (U-Pb /H1100518.6, K-Ar (biotite) /H1100516.0 Ma). On the other hand, a few K-Ar
(biotite) ages are identical within the error limits with their corresponding zircon U-Pbages, such as those of Ichinohashi and Uttsu-dake. This may suggest that the plutonscooled very fast from magmatic to Ar isotope closure temperature of biotite at about250 to 300 °C. In this case, the K-Ar ages can be regarded as the time of magmaemplacement.
In any case, figure 11 suggests that the Miocene peak at about 18 Ma must
represent the most pre-eminent tectonothermal event in Hokkaido. The age spectrum(fig. 11) appears to show a continuous magmatic activity since 52 Ma, but we incline tothink that many of the “intermediate ages” do not represent significant thermal events.
Fig. 7. Chemical characterization of the Hokkaido granitoids. (A) In the Q’-ANOR classification
scheme of Streckeisen and Le Maitre (1979), the granitoids fall in the fields of granodiorite and “granite” ( /H11005
monzogranite and syenogranite). (B) Most granitoid samples are slightly peraluminous and all but one hasA/CNK values (Shand’s index) less than 1.1. Thus, the granitoids are of I-type. (C) In a binary plot ofZr/H11001Nb/H11001Ce/H11001Y vs. FeO*/MgO (Whalen and others, 1987), the granitoids show various degrees of fractional
crystallization. No rocks belong to A-type granite. (D) In a geotectonic classification of granitoids by Pearceand others (1984), all the Hokkaido granitoids fall in the field of volcanic-arc granites.731 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
However, late Miocene granitic emplacements during 12 to 8 Ma have been well
documented (Ishihara and others, 1998; Ishihara, 2007).
Petrogenesis of the Granitoids—Constrained by Geochemical and Sr-Nd-Hf Isotopic
Compositions
The generation of these granitic rocks has been debated. In the northern half of
the chains, two models have been proposed for the generation of the Eocene granite:one is related to ridge subduction and the other to arc magmatism due to change ofsubduction polarity. For the Miocene granites no model has been proposed.
In the southern half, a model for the Eocene granitoids is also related to ridge
subduction, and a model for the Miocene granitoids calls for a mantle upwellingrelated to the opening of the Japan and Kurile basins (Kimura and Kusunoki, 1997;Usuki and others, 2006; Kemp and others, 2007a). In addition, some granitic rocks thatintruded into a regional metamorphic zone are thought to be derived by mixing of abasic-to-intermediate magma and an acid magma of crustal partial melting (Owadaand others, 2006).
The granitoids of the present study were collected from several areas in the central
zone of Hokkaido extending N-S for more than 300 km. Besides, they were formed inthree magmatic stages; thus, any petrogenetic model attempting to relate all the rocksis not realistic. However, it is possible to discuss the general mode of magmageneration based on the geochemical and isotopic characteristics documented inpreceding sections. The validity of proposed models can be tested with the geochemi-cal and isotopic data.
In general, the granitoids show the typical arc magma signatures, including
light-REE enriched rare earth patterns with strong negative Eu anomalies, distinctive
100
10
1
100
10
1
0Rock/Chondrite Rock/Primative mantleAB-1
UK-1ST-1UTT-1UTT-2
AB-1
UK-1
ST-1
UTT-1
UTT-2Older granites (37-48 Ma)
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
RbBaThUKNb Ta La Ce Sr Nd PZr Sm Eu Ti Tb Dy Ho Y Er Yb LuRbBaThUKNbTa La Ce Sr Nd P Zr SmEu Ti Tb Dy Ho Y Er Yb LuLa Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb LuNS-1
SRK-1
ICH-1OTC-2SAH-1
SRK-2
OTC-1Younger granite (ca. 18 Ma)
NS-1
SRK-1
ICH-1
OTC-2SAH-1
SRK-2
OTC-1100
10
1
100
10
1
0AB
CD
Fig. 8. REE distribution patterns and spidergrams for the Hokkaido granitoids and a gabbro. Note that
the older Eocene granitoids and the younger Miocene granitoids are quite similar in the trace elementdistributions.732 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 4
Rb-Sr and Sm-Nd isotopic compositions of granitoids and a gabbro from Hokkaido, Japan
*SRK-2; Use same age with SRK-1 because sampling points are close
*UTT-1; Use same age with UTT-2 because sampling points are closeI(Sr) /H11005(
87Sr/86Sr)sample(t) /H11002(87Rb/86Sr)sample(0) /H11003(e/H9261t/H110021)
εNd(0)/H11005104/H11569((143Nd/144Nd)sample(0) /(143Nd/144Nd)CHUR(0) /H110021)
εNd(T)/H11005104/H11003((143Nd/144Ndsample(0) /H11002147Sm/144Ndsample(0) /H11003(e/H9261147t/H110021))/(143Nd/144NdCHUR(0) /H11002147Sm/144NdCHUR(0) /H11003(e/H9261147t/H110021))/H110021)
fs/H11005fSm/Nd, sample(0) /H11005(147Sm/144Nd)sample(0) /(147Sm/144Nd)CHUR(0) /H110021
TDM-1/H11005(1//H9261147)/H11003loge{1/H11001[(143Nd/144Nd)sample(0) /H11002(143Nd/144Nd)DM(0) ]/[(147Sm/144Nd)sample(0) /H11002(147Sm/144Nd)DM(0) ]}
Constants and parameters used in calculations:/H9261(
87Rb)/H110051.42/H1100310/H110025/Ma
/H9261(147Sm)/H110056.54/H1100310/H110026/Ma (Lugmair and Marti, 1978)
(143Nd/144Nd)CHUR(0) /H110050.512638 (Goldstein and others, 1984)
(147Sm/144Nd)CHUR(0) /H110050.1967 (Jacobsen and Wasserburg, 1980)
(143Nd/144Nd)DM(0) /H110050.51315; (147Sm/144Nd)DM(0) /H110050.2137733 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
negative Ta-Nb-Ti anomalies in the spidergrams, positive whole-rock εNd(t) values ( /H110011
to/H110015), positive zircon εHf(t) values ( /H1100110 to /H1100118), and young whole rock Sm-Nd
model ages (400-1000 Ma), as well as young zircon Lu-Hf model ages ( /H11021400 Ma). Note
that little geochemical and isotopic difference exists between the Eocene and Miocenegranitoids. This suggests that their mode of generation and source rock nature wererather similar but not differentiated by the age factor. The overall isotopic signaturesindicate that the granitoids and the bulk crust of Hokkaido must be quite “juvenile.”However, the positive ε
Nd(t) values ( /H110011t o/H110015) are not so “mantellique”; they are lower
than that of the depleted mantle, thus some amount of older crustal contribution isimplied in the granite petrogenesis. In addition, the large range of zircon Hf isotope
Fig. 9. (A) Whole-rock Sr-Nd isotopic compositions of the analyzed granitoids showing that all eNd(t)
values are positive. The same scenario is found for Cenozoic volcanic rocks in Hokkaido (B). In (B), the fourdata point showing slightly negative e
Nd(t) values are for the “S-type tonalite” xenoliths from the Nozuka-
dake area, Hidaka metamorphic belt (Owada and others, 2006).734 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 5
Hf isotopic data of zircons (in-situ analysis by ICP-MS)735 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 5
(continued)736 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
Table 5
(continued)737 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 5
(continued)
εHf(0)/H1100510000 /H11569{[(176Hf/177Hf)s(0)/(176Hf/177Hf)CHUR(0) /H110021}
εHf(t)/H1100510000 /H11569{[(176Hf/177Hf)s(0)/H11002(176Lu/177Hf)s(0)/H11569(e/H9261/H11569t/H1156910∧6/H110021)]/[(176Hf/177Hf)CHUR(0) /H11002(176Lu/177Hf)CHUR(0) /H11569(e/H9261/H11569t/H1156910∧6/H110021)]/H110021}; t (Ma)
/H9261/H11005/H9261176/H110051.867 /H1156910/H1100211/yr (So ˜derlund and others, 2004)
(176Hf/177Hf)CHUR(0) /H110050.282772 (Blichert-Toft and Albarede, 1997); (176Lu/177Hf)CHUR(0) /H110050.0332 (Blichert-Toft and Albarede, 1997)
fcc/H11005/H110020.5482; fs/H11005(176Lu/177Hf)sample(0) /(176Lu/177Hf)CHUR(0) /H110021
TDM1/H110051//H9261/H11569ln[1/H11001(176Hf/177Hfsample(0) /H11002176Hf/177HfDM(0) )/(176Lu/177Hfsample(0) /H11002176Lu/177HfDM(0) )]
176Hf/177HfDM(0) /H110050.28325 (Griffin and others, 2000);176Lu/177HfDM(0) /H110050.0384 (Griffin and others, 2000)
TDM2/H11005TDM1/H11002(TDM1/H11002t)/H11569(fcc/H11002fs)/(fcc/H11002fDM); fDM/H110050.1566738 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
compositions also attests to the contribution of recycled crust. A crude estimate of the
proportion of juvenile/recycled components is shown in figure 12. The estimate wasdone using a simple two-component mixing calculation, assuming the two end-members to be mantle-derived basaltic rocks ( /H11005mantle) and old continental crust ( /H11005
crust). The mixing proportions for all granitoid samples can be calculated using thefollowing equation:
X
m/H11005Nd c/H11003/H20849/H9255c/H11002/H9255r/H20850//H20851/H9255r/H11003/H20849Nd m/H11002Nd c/H20850/H11002/H20849/H9255m/H11003Nd m/H11002/H9255c/H11003Nd c/H20850/H20852/H11003100
Fig. 10. Zircon Hf isotopic compositions. (A) Zircon grains of the same rock show a significant variation
in Hf isotopic composition, suggesting an open-system behavior during the fractional crystallization in themagma chamber. (B) Consequently, the calculated Lu-Hf model ages also vary within a single rock.Nevertheless, the ensemble of the dataset show relatively young model ages and highly positive ε
Nd(t) values,
indicating the broadly juvenile characteristics of the Hokkaido granitoids.739 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Table 6
Comparison of new zircon ages (in bold type) with K-Ar and Fission-Track age data740 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
where Xm/H11005% mantle-derived juvenile component (represented by basalt); εc,εr,εm/H11005
Nd isotopic compositions of the crust, rock measured, and mantle component,respectively. Nd
c,N dm/H11005Nd concentrations in the crust and mantle components,
respectively. The inset was taken from Jahn (2004), and the parameters used are: εm/H11005
/H110018,εc/H11005/H1100212 (NE China and Inner Mongolia), /H1100230 (Central Mongolia and Transbai-
kalia), /H1100215 (Altai Mountains), /H110024 (Junggar), /H1100215 (Kazakhstan), Ndm/H1100515 ppm, and
Ndc/H1100525 ppm. For the granitoids of Hokkaido, εm/H11005/H1100110 and εc/H11005/H1100210 were used.
We conclude that the juvenile or mantle component in the protoliths of the
Hokkaido granitoids is between 65 to 95 percent. As displayed in the inset of figure 12,among the various tectonic terranes of the CAOB, the Junggar crust, as represented byits Paleozoic granitoids (500-300 Ma), has the most juvenile characteristics, with 60 to100 percent mantle component. Thus, Hokkaido is most comparable with the Junggarterrane regarding the crustal evolution.
In the earlier review of the geology of Hokkaido (section II), we presented that the
central Hokkaido is occupied by the Sorachi-Yezo and Hidaka Belts. The Sorachi-YezoBelt is a Cretaceous-Paleogene forearc basin and accretionary complex; whereas theHidaka Belt is a Paleogene arc complex in the north and the Hidaka metamorphic beltin the south. In a crustal section proposed by Ueda (2005), the likely source region forgranitoid generation, the lower to middle crust, is composed of subducted oceaniccrust (pillowed basalt), seamount, accretionary complex and a sedimentary coversequence. The accretionary complex includes ophiolite m ´elange and other compo-
nents of “ocean plate stratigraphy.” Partial melting of such lithological assemblages
would produce granitic magmas with juvenile isotopic characteristics as shown in thisstudy.
Fig. 11. Summary of the available age data (new and literature) for the intrusive rocks from Hokkaido.
Literature data source: Arita and others (1993, 2001), Honma and Fujimaki (1997), Ishihara and Terashima(1985), Ishihara and others (1998), Kawakami and others (2006), Kawano and Ueda (1967), Kemp andothers (2007a), Kimbrough and others (1994), Koshimizu and others (1988), Koshimizu and Kim (1986),Kubo and others (1984), Maeda and others (1990), Nakagawa (1992), Okamoto and Honma (1983),Okamura and others (2003), Ono (2002), Owada and others (1997, 2006), Saheki and others (1995),Shibata (1968), Shibata and others (1975), Shibata and Ishihara (1979, 1981), Zeniya and others (1996).741 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
Ishihara and others (1998) noted that a distinctive feature of the late Cenozoic
plutonism in the north-central Hokkaido is that granitoids and gabbroids occur in anequal amount (120 vs 110 km
2) and are closely associated. The granitoids tend to occur
in the whole region, but the gabbroids are restricted to the western edge of the Hidakabelt. Since the emplacement times of granitoids and gabbroids are similar ( ca.18 Ma),
the plutonism is considered as bimodal in character. Moreover, bimodal magmatism isgenerally known to occur in an extensional tectonic regime, therefore, many modelsare in favor of the generation of the Hokkaido Miocene granitoids in a back-arc setting.On the other hand, the Eocene granotoid magmas were probably produced by meltingof subducted accretionary complexes in supra-subduction zones. The accretionarycomplexes were likely dominated by juvenile or mantle-derived lithological assem-blages as argued from the Sr-Nd-Hf isotopic signatures.
Juvenile Crustal Growth and Comparison with Other Parts of Japan
The formation of the Japanese Islands has been taken as a standard model for
accretionary orogeny. It was proposed that the most important cause of the orogeny isthe subduction of oceanic ridge, by which the continental mass increases through thetransfer of granitic melt from the subducting oceanic crust to the orogenic belt(Maruyama, 1997). Sengor and Natal’in (1996) named the orogenic complexes“Nipponides,” consisting predominantly of Permian to Recent subduction-accretioncomplexes with few fragments of old continental crust, and further pointed out theresemblance in orogenic style between Japan and the Central Asian Orogenic Belt(CAOB). Consequently, the Japanese Islands are essentially built up by juvenile crust.
Fig. 12. Proportion of the mantle (juvenile) component involved in the generation of the Hokkaido
granitoids. Assumption in the estimate for the Hokkaido rocks: Nd concentrations used for the mantle-derived juvenile component (“m”) and recycled crustal component (“c”) are 15 and 25 ppm, respectively.The Nd isotopic compositions used for ε
Ndm/H11005/H1100110 and εNdc/H11005/H1100210.742 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
However, based on the available Sr-Nd isotopic data, Jahn (2010) showed that a large
proportion of the granitoids of SW Japan have Proterozoic Sm-Nd model ages, highinitial
87Sr/86Sr ratios and negative εNd(T) values (fig. 13). These isotopic data are in
strong contrast with those of two celebrated accretionary orogens: the Central AsianOrogenic Belt and Arabian-Nubian Shield, but are quite comparable with thoseobserved in SE China and Taiwan, or in classical collisional orogens in the EuropeanHercynides and Caledonides (Jahn, 2004). This raises questions about the bulkcomposition or type of material accreted in accretionary complexes, and negates thehypothesis that the Nipponides contains very few fragments of older continental crust(Sengor and Natal’in, 1996). Jahn (2010) concluded that the subduction-accretioncomplexes in SW Japan were composed in significant amount of recycled continentalcrust, probably of Proterozoic age. The scenario is comparable with that in Taiwan.
However, further research has revealed that the real juvenile crust was produced
in other parts of Japan, rather than in the best-studied SW Japan. As demonstrated inpreceding sections, Hokkaido as a whole provides an excellent example of juvenilecrustal addition to the global continental crust. Figure 14 illustrates the essential Nd-Srisotopic plots for the granitoids from NE Japan (Hokkaido included). The data ofgranitoids ( sensu lato ) from the Kitakami and Abukuma Mountains and the Niigata
area are shown for comparison with that of the granites and rhyolites from Hokkaido.Note that many of them are new and unpublished data.
Figure 14 shows that the majority of the data points have positive ε
Nd(T) values
and initial87Sr/86Sr ratios of /H110170.7055. The data of Abukuma granitoids straddle the
εNd(T) zero-line but generally form a negative correlation with the rest of the data
points. By contrast, the data of the Niigata granitoids (shown in blue squares) seem to
Fig. 13. Nd-Sr isotopic plots of granitoids from SW Japan (original data are from Jahn, 2010 and the
cited references).743 zircon geochronology, Sr-Nd-Hf isotopic and geochemical analyses
form a separate group distinguished from the rest but is comparable with the data
array of SW Japan (fig. 13). Note that the data arrays in figures 13 and 14 are highlycontrasted. The juvenile-crust-dominated features in NE Japan are replaced by therecycled-crust-dominated characteristics of SW Japan. This indicates that the architec-ture and crustal evolution of the two major parts of the Japanese Islands are quitedistinguished. Note that the isotopic data displayed in figures 13 and 14 [ ε
Nd(T) vs
(87Sr/86Sr)o] involve rocks of different ages, so their display on the same plane is not
strictly valid. In theory, we should have adjusted all the data points to the same timeline. However, the adjusted vectors for a time difference less than 100 Ma would be toosmall to be detected in the figures, so the calculated initial ratios were plotted directlyon the same plane.
The comparable isotopic compositions and the occurrence of the Niigata grani-
toids to the west of the Tanakura Tectonic Line lend support to the idea that thetectonic boundary or suture zone between NE and SW Japan before the Cenozoic ismore logically represented by the Tanakura Tectonic Line, but not the Itoigawa-Shizuoka Fault.
conclusions
The present study leads to the following conclusions:1. Zircon U-Pb geochronology revealed three distinct periods of granitoid emplace-
ment in central Hokkaido, at 45, 37 and 18 Ma.
2. Geochemical analyses show that the granitoids comprise granodiorite, monzo-
granite and syenogranite; they are weakly peraluminous and possess volcanic arccharacteristics. They are not S-type granites. All the granitoids, regardless of their ages
Fig. 14. Nd-Sr isotopic plots of granitoids from NE Japan (Hokkaido included). Individual source of
data are not listed herein, but the original data-set used in this compilation is available upon request.744 B.-M. Jahn & others—Cenozoic granitoids in Hokkaido (Japan): Constraints from
‘(Eocene or Miocene), have similar REE patterns and spidergrams, typical of Phanero-
zoic granitoids.
3. Whole-rock Sr-Nd and zircon Hf isotopic data indicate that the granitoids are
quite juvenile and likely generated by partial melting of sources dominated bymantle-derived rocks, and in matured arc settings. Recycled ancient crustal rocks arenot a significant component in the source regions of the granitoids, and probably theentire Hokkaido crust.
4. The literature data show that the Miocene volcanic rocks (rhyolites, dacites and
andesites) from Hokkaido possess Sr-Nd isotopic characteristics comparable with thegranitoids, hence the plutonic granitoid and volcanic felsic magmas were probablyderived from similar juvenile sources.
5. The crustal development of Hokkaido is most comparable with that of the
Junggar Terrane of the Central Asian Orogenic Belt. Together with the Arabian-Nubian Shield, they document the best examples of juvenile crust growth.
acknowledgments
We have benefitted from frequent communications with Dr. Shunso Ishihara
(Geological Survey of Japan/AIST) and Dr. Kazu Okamoto (Saitama University).Shoujie Liu and Jeremy Wu helped in preparation of some figures presented herein.Professor Dunyi Liu and the Beijing SHRIMP Center provided a temporary shelter toBMJ in May 2013 for preparation of the manuscript. Drs. Ishihara, Ueda and BillGriffin critically reviewed this paper and helped ameliorate the manuscript. Dr. Uedais particularly thanked for improving the presentation of geology and tectonics ofHokkaido. Bor-ming Jahn acknowledges the support of NSC-Taiwan (NSC100-2116-M-002-024, NSC100-2923-M-002-010 and NSC101-2116-M-002-003).
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Jahn (2014) - Generation of Cenozoic granitoids in Hokkaido.txt
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Geology and fluid chemistry of the Fushime
geothermal field, Kyushu, Japan
Hiroaki Okadaa,*, Yoshio Yasudaa, Masahiko Yagib,
Kunio Kaia
aJapan Petroleum Exploration Co., Ltd, 2-2-20 Higashi-Shinagawa, Shinagawa-ku,
Tokyo, 140-0002, Japan
bJapan Petroleum Exploration Co., Ltd, Research Center, 1-2-1 Hamada, Mihama-ku,
Chiba, 261-0025, Japan
Received 26 January 1998; accepted 7 July 1999
Abstract
The Fushime geothermal field is located in a depression close to the coast line. The
system is characterized by very high reservoir temperature (>350 8C), and a high salinity
production fluid. Geological analysis shows that the main reservoir in this field occurs in a
fractured zone developed around a dacite intrusion located in the center of the field. High
permeability zones recognized by drilling data are found to be associated with fault zones.One of these zones is clearly associated with a NW–SE trending andesite dike swarm which
was encountered in some wells.
Alteration in the system can be divided into four zones, in order of increasing
temperature, based on calcium–magnesium aluminosilicate mineral assemblages: i.e., thesmectite, transition, chlorite and epidote zones. The feed zone is located in the chlorite and
epidote zones, which can be further divided into three sub-zones according to their
potassium or sodium aluminosilicate mineralogy, from the center of the discharge zone: K-feldspar–quartz, sericite–quartz, and albite–chlorite zones.
Chloride concentration of the sea-water is 19,800 mg/l, and Br/Cl mole ratio is 1.55.
Based on geochemical information, the reservoir chloride concentration of this field rangesfrom 11,600 to 22,000 mg/kg. The Clres (Cl in reservoir), Br/Cl ratios and stable isotope
data indicate that the Fushime geothermal fluid originated from sea-water and is diluted by
ground water during its ascent. Some fluids produced from geothermal wells show low pH
0375-6505/00/$20.00 72000 CNR. Published by Elsevier Science Ltd. All rights reserved.
PII: S0375-6505(99)00063-2Geothermics 29 (2000) 279–311
* Corresponding author. Tel.: +81-3-5461-7334; fax: +81-3-5461-7394.
E-mail address: okada@gtm.japex.co.jp (H. Okada).
(about 4). It is thought that sulfide mineral (PbS, ZnS) precipitation during production
produces this acidic fluid. 72000 CNR. Published by Elsevier Science Ltd. All rights
reserved.
Keywords: Geothermal system; Yamagawa power station; Fushime geothermal field; Japan
1. Introduction
The Fushime geothermal field is located in the southernmost part of Satsuma
Peninsula (called the ‘‘Satsunan’’ area) in southern Kyushu, Japan (Fig. 1). Theinitial geothermal investigation of the Satsunan area was carried out by theJapanese government. An exploratory well (SA-1) drilled in the Fushime area in
1975 recorded temperatures exceeding 230 8C at a depth of 500 m. NEDO (New
Energy and Industrial Technology Development Organization) conducted a seriesof surveys in the northwestern part of the Satsunan area during a geothermalsurvey in 1982 and 1983. This program included the drilling of six wells ranging indepth from 1000 to 1700 m.
Japan Petroleum Exploration Co., Ltd (JAPEX) has carried out geothermal
exploration in this field since 1977. JAPEX Geothermal Kyushu Ltd (JGK), asubsidiary of JAPEX, has continued exploitation at this field since 1988. As a
result, the Yamagawa power station began operation here in March 1995, with 30
MWe of geothermal power generation.
The Fushime field is situated in an agricultural area at an elevation of about
40 m, near the coastline facing the East China Sea. The entire site of the powerplant, including the production/reinjection facilities, occupies an area of approx.160,000 m
2only, because most wells are directional (Fig. 2). At present, 12
production wells ranging from about 1500 to 2500 m in drilled depth, and fourreinjection wells ranging from about 800 to 1300 m in drilled depth, are being
utilized. The total discharge of steam is about 225 t/h, separated at a pressure of
about 10 kg/cm
2gauge. Hot water of about 350 t/h is reinjected into wells at the
separated pressure to avoid silica precipitation in pipes. The field is in closeproximity to the coast; production fluid is rather saline, suggesting its origin fromsea water.
During the exploration and exploitation phases, a variety of geological,
geophysical and geochemical surveys and analyses has been conducted, as well asthe drilling of over 20 wells.
This paper describes the geologic structure and its role within the geothermal
system, and the fluid chemical characteristics of the Fushime field.
2. Regional geology
In southern Kyushu, four huge calderas, Kikai, Ata, Aira and Kakuto-H. Okada et al. / Geothermics 29 (2000) 279–311 280
Kobayashi occur at 50–60 km intervals in a SSW–NNE direction along
Kagoshima Bay (Fig. 1). The Onkadobira fault extends in a SW–NE directionacross the Satsuma peninsula and is a major structure separating the Satsunanarea from the rest of the peninsula. The area also corresponds to the western partof the large Ata caldera (Matumoto, 1943), with the Onkadobira fault defining thenorthwestern margin of the caldera. The collapse of the Ata caldera is thought tohave accompanied the multiple eruptions that produced the voluminouspyroclastic deposits between 105–110 ka (Matsumoto and Ui, 1997).
The generalized stratigraphy of the Satsunan area is shown in Table 1,
Fig. 1. Location of Satsunan area, Fushime geothermal field, and the Yamagawa geothermal power
station. The field is located within the Kagoshima graben (Tsuyuki, 1969), a volcano-tectonic
depression, and the Ata caldera (Matumoto, 1943).H. Okada et al. / Geothermics 29 (2000) 279–311 281
comprising Quaternary pyroclastic rocks and lavas overlying Tertiary formations.
A simplified geologic map is shown in Fig. 3. The basement in this area isconsidered to be composed of Mesozoic–Paleogene marine mudstones andsandstones of the Shimanto supergroup, and a granitic pluton with K-Ar ages of14–22 Ma (Aramaki and Ui, 1966) which intruded into the Shimanto supergroup.Although these basement rocks do not outcrop in this area, granitic rock wasencountered in NEDO’s well ID-2, located on the north shore of Lake Ikeda (Fig.3), at a depth of 452 m (NEDO, 1986). Most of the pre-Ata rocks, including theYamagawa formation and Nansatsu group, occur on the upthrown side of theOnkadobira fault.
The topographic features are dominated by post-Ata lava domes and flows of
the central cone massif within the Ata caldera, extending between Yamagawa bayand the Onkadobira fault. The central cone eruptive sequence comprised multipleandesitic lava and minor pyroclastic eruptions spread over an interval of at least20,000 years. Well preserved southern lava domes include Kasegidake,Tsujinodake and Kogawa.
The most recent eruptions in the central cone area consisted of the Ikeda
sequence about 5000–5500 years B.P. (Naruo, 1983), forming the 4 3k m
Fig. 2. View looking southeast of Yamagawa geothermal power station, with the Takeyama peaks and
the East China sea in the background.H. Okada et al. / Geothermics 29 (2000) 279–311 282
Table 1
Stratigraphy of the Satsunan area
EPOCH Stratigraphy Major volcanoes Rock type Age
Holocene Post-Ikeda volcanic rocks Kaimondake,
NabeshimadakeAndesite lava,
pyroclastics, basalt,
daciteKaimondake;
14C 3620 2140 year B.P. (Furukawa and Nakamura,
1969),14C 4040 2120 year B.P. (Ishikawa et al., 1979),
1000–4000 year B.P. (Nakamura, 1971)
Ikeda pyroclastic rocks Ikeda caldera Pumice fall, pumice
flow, pyroclastic rocks14C 4640 280 year,
T.L. 4550 year,
Ikeda volcano 5000–5500 year B.P. (Naruo, 1983)
Pleistocene Post-Ata volcanic rocks Washiodake,
Ikezoko lava,
Kiyomidake,Takeyama,Yamagawa lava, etc.Andesite lava,
pyroclastic rocksKiyomidake; F.T. 25,000 year B.P.,
Takeyama; F.T. 26,000 year B.P.,
Yamagawa lava; F.T. 27,000 year B.P.(Kamitani et al., 1976)
Ata pyroclastic rocks Ata caldera Pyroclastic flow,
welded tu/C128105–110 ka (Matsumoto and Ui, 1997)
Pre-Ata volcanic rocks Irino lava,
Kohama lava,
Yahazudake,Takaeyama, etc.Andesite lava, dacite
lavaIrino lava F.T. 0.17 20.04 Ma,
Kohama lava K-Ar 0.8 20.6 Ma,
Yahazudake F.T. 1.7 20.3 Ma,
Takaeyama F.T. 2.1. 20.3 Ma
Early Pleistocene Yamagawa formation Tu/C128, tu/C128 breccia N22 zone
a, N20–N21 zonea
0
Miocene Nansatsu group Tu/C128, tu/C128 breccia,
andesite, rhyolite
Early Miocene
0Basement Rock
IntrusionGranitic rocks Granite (southern Osumi),
K-Ar 14 21 Ma, 21 21 Ma, 22 Ma (Aramaki and Ui,
1966),
Granite (Satsuma) F.T. 14–15 Ma (Miyachi, 1990)
Palaeogene
0
MesozoicShimanto supergroup Marine sediments
(mudstone, sandstone)
aZoning of Blow (1969).H. Okada et al. / Geothermics 29 (2000) 279–311 283
Fig. 3. Simplified geological map with gravity of the Satsunan area. Most pre-Ata rocks are located on the upthrown side of the Onkadobira fault. On the
downthrown side, young, post-Ata volcanoes form a central massif. The Fushime geothermal field lies within the gravity low.H. Okada et al. / Geothermics 29 (2000) 279–311 284
Krakatau-type caldera (Ui, 1967), and a line of explosion craters extending 5 km
southeast to Yamagawa bay (Fig. 4). Ejecta from other explosion craters atNarikawa, Unagi, Ikezoko and Matsugakubo occupy the same stratigraphic
position within the Ikeda deposits. These explosion breccia deposits appear to
have been produced by the intrusion of dikes along fissures extending between theIkeda caldera and Yamagawa bay.
The Kaimondake cone grew to 922 m a.s.l. accompanied by a scoria fall over
the entire Satsunan area, and was the major event after the Ikeda eruptions.Kaimondake comprises a basal stratovolcano with a small central lava dome ontop. The most recent eruptions of Kaimondake were recorded in 884 and 885A.D. (Fujino and Kobayashi, 1997).
Fig. 4 shows the collapse features and distribution of major thermal
manifestations in the Satsunan area. There are numerous high temperature hotsprings and several fumaroles. Ibusuki is famous as a hot spring resort, and thereare many other hot spring areas such as Unagi, Narikawa, Fushime andNagasakibana. Hot springs in these areas are used directly for bathing, heatinglocal greenhouses and fish hatcheries. Most of the fumaroles are distributed inareas of high relief, while the hot springs are located in flatlands. Most of thesehigh temperature hot springs are saline except for Unagi.
An analysis of trends of volcanic vent locations indicates some NW–WNW
lineations. The most clearly defined lineation is that passing through the summitsof Tsujinodake, Kogawa domes and the unexposed dome to the southeast. Thisnorthwest trend continues directly through the summits of the Takeyama peaks tothe volcanic neck of Matagoshi Island. Takeyama is a topographic high along thecoast just south of the Fushime area, and is a northwest trending ridge with steepsided peaks on its northwest and southeast margins (Figs. 2 and 3). Both peaksare underlain by a pyroxene andesite. This northwest trend is interpreted as
marking an important buried fault along which magmas have been extruded. The
Ikeda-age craters of Matsugakubo and Ikezoko also define a northwest trendwhich continues through small craters or collapse features to the fumarolic areaon the northeastern shore of Lake Unagi. O/C128set to the south, there is a parallellineation through the main Unagi, Narikawa and Yamagawa explosion craters.These northwest trends, normal to the Onkadobira fault, appear to be a majorfeature of volcanism in the Satsunan area.
3. Geology of the Fushime field
3.1. Stratigraphy
Subsurface geology has been investigated by drill cuttings collected every 5 m
and a number of spot cores from over 20 wells ranging from 755 to 2605 m indepth (average drilled depth is about 2000 m). The stratigraphy encountered inFushime wells (Table 2) can be divided into six units (Yoshimura et al., 1985).H. Okada et al. / Geothermics 29 (2000) 279–311 285
Fig. 4. Topographic features and surface manifestations, including the result of a resistivity survey. Collapse features, eruption centers of youn g volcanoes,
and the distribution of hot springs and fumaroles are shown. Most hot springs discharge from shallow wells. Fushime geothermal field lies within theresistivity low.H. Okada et al. / Geothermics 29 (2000) 279–311 286
Table 2
Stratigraphy of Fushime wells
EPOCH Stratigraphy RemarksHolocene Kaimondake scoria Scoria fall from Kaimondake. It covers the area widely.
Ikeda pyroclastic rocks Mainly composed of pumice flow deposits.
Late Pleistocene Takeyama andesite Andesite lava related to Takeyama volcano.
Fushime silt Occurs in the Fushime wells about 120 m below the surface, and contains a lot
of foraminifera that are correlative with the Shiroyama formation (33,000–110,000 year ) in Kagoshima.
Early Pleistocene
0Yamagawa formation Mainly composed of dacitic tu/C128 and tu/C128 breccia with occasional thin beds of
siltstone. Foraminifera data indicate N21, N22 zones
a.
Late Pliocene Fushime welded tu/C128 Strongly welded tu/C128 occurs at the base of this formation locally.
Pliocene Nansatsu group
Upper Nansatsu
formationFormation is altered dacitic tu/C128 and tu/C128 breccia, which are commonly
silicified. Foraminifera data indicate N17–N21 zonea.
Middle NansatsuformationFormation is composed of andesitic lava and pyroclastics, and occurred in the
wells drilled in the northern and southern part of the Fushime field.
0
Miocene Lower Nansatsu form Formation is composed of altered dacitic tu/C128.
aZoning of Blow (1969).H. Okada et al. / Geothermics 29 (2000) 279–311 287
From the surface, they are: Kaimondake scoria, Ikeda pyroclastic rocks,
Takeyama andesite, Fushime silt, Yamagawa formation, and Nansatsu group.
The Kaimondake scoria is a widespread black, reddish-black scoria fall with
volcanic lapilli from the Kaimondake volcano. The Ikeda pyroclastic ejecta is
composed of pumice flow deposits and its reworked sand and gravel depositsrelated to the formation of the Ikeda caldera. Thickness in the field is about100 m. The Takeyama andesite is composed of dark gray-to-black andesite lavawhich was derived from the Takeyama volcano. The Fushime silt consists of alight-to-dark gray tu/C128aceous silt and yields abundant shell fragments andforaminifera such as Elphidium advenum, Pseudononion japonica and Operculina
ammonoides . It correlates with the Shiroyama formation (30–110 ka) in
Kagoshima City on the basis of foraminiferal fauna (Yoshimura et al., 1985).
The Yamagawa formation is mainly a porous dacitic pyroclastic unit that
occurs from about 150 m to about 1500 m depth with some thin siltstone beds.Foraminifera from these siltstone beds indicate a correlation with the Globigerina
pachyderma (sinistral)/Gna. incompta zone of Maiya (1973), which corresponds
approximately to the N22 zone defined by Blow (1969), according to Maiya(1973). In addition, foraminifera indicative of Blow’s (1969) N21 zone are alsorecorded from the formation. This formation is thus considered to be late
Pliocene-to-early Pleistocene in age. At the base of this formation, a thick,
strongly welded tu/C128 occurs locally.
The Nansatsu group encountered in this field is divided into three formations
based on lithology. The Upper Nansatsu formation consists mainly of daciticpyroclastic rock with some thin siltstone beds. Foraminifera also occur in somesiltstone beds. Foraminifera such as Globigerina bulloides indicate the N17–21
zone; thus this formation is considered to be late Miocene-to-Pliocene in age. TheMiddle Nansatsu formation consists mainly of pyroxene andesite lava and
andesitic pyroclastic rocks. The Lower Nansatsu formation is composed of
strongly silicified dacitic tu/C128s.
In addition, there are also some intrusions present. The largest is a dacite in the
central part of the field, which intruded the Upper Nansatsu formation. It is aporphyritic, felsitic and sometimes spherulitic hypersthene dacite. Another clearlydistinguished intrusion is an andesite dike swarm. These andesite dikes have beenencountered in three directional wells drilled to the south.
3.2. Structure
The Fushime field lies in a depression that is clearly identified from a gravity
low (Fig. 3) and resistivity low (Fig. 4). This depression has been also confirmed
by the geologic correlation of wells. The deepest well in the Fushime field, well 3,was drilled to a depth of 2605 m. No well in the field has, however, reached thebasement rock of either the Shimanto supergroup or granitic rock. Using thedepth of granitic rock encountered in NEDO’s well ID-2 as a control point, thedepth of the basement in the Fushime field is estimated at about 3.5 km, based ona gravity analysis (Yoshimura et al., 1985).H. Okada et al. / Geothermics 29 (2000) 279–311 288
As this area is situated in a farm area and provides few outcrops, the
subsurface geologic structure was constructed mainly on the basis of geologiccorrelations between wells over an area of approximately 4 km
2, plus geophysical
survey results.
Some faults are inferred to occur inside the depression and are predominantly
downthrown towards the center of the depression (Figs. 5 and 6). In addition tothese faults, a northwest structural trend that passes through Takeyama is inferredto be a fault. This trend is also observed as a high resistivity axis by amagnetotelluric (MT) survey, possibly caused by a dike near Takeyama. In wells7, 15 and 23, many andesite units between about 700 and 1200 m have beenconfirmed by cuttings, well logs and drilling data (Fig. 7). These andesite units are
not encountered in neighboring wells, and the depths of the andesite units di/C128er
from well to well. These andesites are clearly associated with the trend passingthrough Takeyama and Tsujinodake volcanoes (Fig. 5), suggesting that they arepart of a dike swarm related to the trend [Takeyama–Tsujinodake line (TTL)].The dike swarm zone is about 100 m wide and extends in a NW–SE direction, themajor trend in the Satsunan area. The trend of the dikes may have been normalto the minimum compressional axis (Nakamura, 1977), suggesting that the leasthorizontal principal stress in this field is oriented NE–SW, approximately normal
to the direction of the monogenic volcanoes and craters alignments in the
Satsunan area.
3.3. Alteration mineralogy
Microscopy, X-ray di/C128raction (XRD) and X-ray fluorescence analysis revealed
the alteration mineralogy in the Fushime field. The alteration can be divided intofour zones on the basis of calcium–magnesium aluminosilicate mineralogy, i.e. thesmectite, transitional, chlorite and epidote zones in order of increasingtemperature (Fig. 8). The smectite zone occurs where the temperature is below1608C; clinoptilolite and a-cristobalite or quartz are common. The transitional
zone is characterized by the occurrence of smectite/chlorite and smectite/sericite
mixed-layered minerals commonly coexisting with quartz, mordenite, andanalcime. The chlorite zone occurs at temperature from 180–200 to 250 8C,
coexisting with quartz, albite and wairakite. The epidote zone is defined by thefirst appearance of radial or fan-shaped epidote, distinguishing it from Tertiaryalteration, at temperatures above 250 8C. Quartz, chlorite and albite are commonly
associated with epidote (Yagi, 1990).
The production zones are usually located in the chlorite and epidote zones.
These zones are characterized by an assemblage of K-bearing minerals that result
in a high concentration of potassium in rocks. Moreover, these zones can besubdivided into three sub-zones according to the assemblage of potassium orsodium aluminosilicate minerals, K-feldspar–quartz, sericite–quartz, and albite–chlorite (Yagi and Kai, 1990). The K-feldspar–quartz assemblage occurs in thecenter of the discharge zone at temperatures above 220 8C. Interstitial or matrix-
replacing K-feldspar is present in association with quartz overgrowths. Matrix-H. Okada et al. / Geothermics 29 (2000) 279–311 289
Fig. 5. Geological structure of the Fushime geothermal field with magnetotelluric (MT) survey results.
Inferred faults and dikes were confirmed by well data.H. Okada et al. / Geothermics 29 (2000) 279–311 290
replacing albite, hydrothermal epidote in radial or fan shapes, and pore-filling
anhydrite usually coexist with K-feldspar. The zone of sericite–quartz assemblageusually surrounds the K-feldspar–quartz zone, and minute crystals of sericite fill a
matrix of authigenic quartz. Matrix-replacing albite, wairakite, and chlorite are
common minerals in this assemblage, but K-feldspar is absent. The zone of thealbite–chlorite assemblage occurs at the outermost portion of the hydrothermal
Fig. 6. North–South geological cross section with temperature distribution of the Fushime geothermal
field. Section line is shown in Fig. 5.H. Okada et al. / Geothermics 29 (2000) 279–311 291
Fig. 7. Geological columns of three directional wells intersecting the Takeyama–Tsujinodake line
(TTL). Location of circulation losses and rates are also shown. Andesite dikes and lost circulation areconcentrated around the TTL.H. Okada et al. / Geothermics 29 (2000) 279–311 292
reservoir. Matrix-replacing albite and chlorite coexist with quartz in this
assemblage, and pore-filling calcite is also present.
3.4. Fluid inclusion thermometry
The homogenization temperatures (Th) of fluid inclusions in cores and cutting
samples were measured (Fig. 9). These inclusions are all hosted by anhydrite,which occurs widely in the geothermal system. The most common inclusions areliquid-rich with uniform liquid–vapor ratios. Vapor-rich inclusions, with varyingvapor–liquid ratios, are also present in the deep reservoir, below 1600 m, and alsoin the shallow reservoir at around 500 m depth. The maximum Th values below
1600 m in most wells are similar to the measured subsurface temperatures. The
values of Th vary widely between the depth of 700 and 1100 m, which iscoincident with a measured temperature inversion zone. The range of Th valuesfrom well 17 span the interval between the boiling curve and the cooler presentmeasured temperature, supporting a recent cooling in this zone, based on thearguments presented by Taguchi et al. (1985).
The crushing method (Sasada et al., 1986) was applied to the fluid inclusions in
this field. The result shows that the CO
2content of the fluid inclusions ranges
from 0.05 to 0.15 mol%. Thus, the maximum decrease of freezing point (Tm)
caused by CO 2is 0.18C. The Tm values in the deep reservoir range from ÿ1.9 to
Fig. 8. Temperature ranges for the occurrence of hydrothermal minerals in the Fushime geothermal
system. Solid lines denote common occurrence, and dashed lines denote rare occurrence. Modified fromYagi (1990).H. Okada et al. / Geothermics 29 (2000) 279–311 293
ÿ2.1, equal to a Cl equivalent of 18,000 to 20,000 ppm after subtracting the CO 2
e/C128ect. This is close to the chloride concentration of sea water. Thus, the chloride
concentration of the fluid inclusions in the deep reservoir suggests that the origin
of the reservoir fluids is sea water.
4. Fluid chemistry
4.1. Chemical composition of hot water
The geothermal fluid collected from eleven production wells was analyzed to
determine the origin of the fluid and to estimate the reservoir characteristics
(Table 3). The pH values measured after quenching to 25 8C in the field range
Fig. 9. Homogenization temperatures of vapor–liquid fluid inclusions in the Fushime geothermal field;
measured temperature and the boiling curve of water are also drawn. Solid line denotes a boiling curve
of pure water, and dashed line denotes that of water with 19,000 ppm Cl.H. Okada et al. / Geothermics 29 (2000) 279–311 294
Table 3
The analytical results of fluids sampled from 11 production wells in the Fushime fielda
Well
7 9 10 11 12 16 17 18 19 20 22 Standard sea
water
Sampling date 30/1/90 31/1/90 31/1/90 30/1/90 22/1/90 22/1/90 29/1/90 29/1/90 15/7/94 25/8/94 12/7/96
Total discharge enthalpy (kJ/kg) 2123 1496 1509 1162 1526 2362 2324 1371 1998 1522 n.a. –
Temperature measured by
temperature logging ( 8C)> 300 290–320 280 250 > 320 330–340 > 330 > 300 n.m. n.m. n.m. –
Gas sample
Sampling press (kg/cm3). 5.9 4.0 12.4 7.5 3.4 4.5 3.7 4.2 4.0 8.7 n.a. –
CO2(mmol/kg) 17.3 127 284 62.9 50.0 25.7 33.1 83.5 76.0 109 n.a. –
H2S 11.3 9.46 11.1 2.49 15.3 16.2 16.7 4.14 19.5 17.2 n.a. –
H2 1.29 1.08 1.69 0.59 1.78 1.22 1.50 0.96 1.56 1.56 n.a. –
N2 0.95 2.49 1.95 3.73 1.96 0.82 0.86 2.36 1.19 3.39 n.a. –
CH 4 0.25 0.55 0.40 0.66 0.27 0.12 0.20 0.30 0.28 0.58 n.a. –
NH 3 < 0.17 0.51 0.79 0.46 – < 0.17 < 0.17 0.25 1.15 1.00 n.a. –
Water sample (collected at atmospheric
pressure)
pH (25 8C) 4.15 6.73 6.94 7.23 6.51 4.88 3.99 7.06 5.77 7.41 4.50 –
Na (mg/kg) 19,100 10,100 13,400 9100 13,400 18,100 15,100 11,500 11,200 8400 10,400 11,000
K 5340 2090 3120 1520 4130 5200 4850 2480 3380 1210 2180 413
Ca 2310 1270 2500 1290 1500 2000 1490 1970 1240 1150 1240 423
Mg 9.8 2.9 2.6 1.4 2.1 13.3 14.0 2.6 5.7 2.3 3.2 1325
Cl 38,600 20,700 27,100 18,400 27,800 35,600 31,100 24,100 24,300 16,100 20,400 19,800B 90.0 46.6 51.6 36.5 74.2 112 133 46.3 83.8 30.7 50.0 4.5
SO
4 49.8 22.5 40.8 31.4 22.1 36.5 54.1 28.2 69.6 44.4 14.6 2720
T–Fe 31.8 0.866 0.281 0.111 0.620 24.4 69.6 0.426 5.78 0.046 8.04 2
T–SiO 2 922 742 595 571 634 922 1274 739 984 584 809 –
dO(-)bÿ1.2ÿ1.3ÿ1.8ÿ1.6ÿ0.4ÿ0.6ÿ0.8ÿ1.1ÿ1.2ÿ2.0 n.a. ÿ0.1c
dD(-)bÿ7.5ÿ8.1ÿ6.4ÿ10.1ÿ3.6ÿ2.4ÿ4.2ÿ4.5ÿ9.1ÿ15.1 n.a. –
Na–K–Ca temperature ( 8C) 307 272 280 253 311 310 319 273 307 242 274 –
SiO 2temperature 273 255 238 235 243 273 302 255 279 237 262 –
Clres (mg/kg) 22,000 13,700 17,300 12,800 16,000 20,000 16,800 15,800 13,800 11,600 13,300 –
Na/Cl10 7.63 7.52 7.61 7.63 7.43 7.83 7.49 7.35 7.10 8.04 7.86 1.03
K/Cl100 12.5 9.2 10.4 7.5 13.5 13.2 14.1 9.3 12.6 6.8 9.7 1.89
Ca/Cl100 5.29 5.43 8.16 6.20 4.77 4.97 4.24 7.22 4.51 6.31 5.38 1.89
Br/Cl1000 1.57 1.47d1.63 1.54 1.55 1.55 1.48e1.55 – – – 1.55
aClres: Cl concentration recalculated to the reservoir condition using the Na–K–Ca geothermometer. n.a.: not analyzed; n.m.: not measured. Each rat io
represents mole ratio. SiO 2temp.=(1522/5.75ÿlogc)ÿ273.15.
bThe values recalculated to the reservoir condition.
cLocal sea water value.
dSampled on May 11, 1994.
eSampled on November 2, 1994.H. Okada et al. / Geothermics 29 (2000) 279–311 295
from 3.99 to 7.41. The fluids separated at atmospheric pressure have a high
chloride concentration, ranging from 16,100 to 38,600 mg/kg.
Chemical geothermometers, such as silica, alkali, gas and isotope, were
calculated to evaluate the reservoir temperature. The silica geothermometer isgenerally used to calculate reservoir temperatures of less than 250 8C. However,
the measured temperatures in several wells exceed 300 8C in the Fushime field. The
temperature calculated by the silica geothermometer is, moreover, lower thantemperature logging data. On the other hand, the Na–K–Ca geothermometertemperature is in good agreement with the temperature of the principal feed zonesas measured by logging (Table 3). The reservoir chloride concentration (Clres) was
therefore recalculated for a liquid at the Na–K–Ca geothermometer temperature,
and these values range from 11,600 to 22,000 mg/kg. These Cl concentrations arelower than or equal to that of sea water, except for two wells. The Br/Cl ratios ofthe fluid samples are also similar to that of sea water. The Clres, Br/Cl ratio andisotope results, to be discussed later, suggest that the Fushime geothermal fluidoriginated from sea water.
Fig. 10 shows the relationship between the Clres and the calculated enthalpy at
the Na–K–Ca geothermometer temperature. As shown in the figure, it is clearthat two di/C128erent types of hot springs exist in this area: one has high Clconcentration and high temperature, the other has low Cl and relatively lowtemperature. The former spring type originates from deep reservoir fluid mixedwith ground water, and the latter may originate from steam-heated water, asdemonstrated by Hedenquist (1990) at Broadlands, New Zealand. Since themaximum measured temperature exceeds 350 8C at a depth of 2000 m in several
wells, sea water might be altered during such heating. The graph indicates thatafter sea water was heated to 350 8C, boiling and/or dilution occurred in the
reservoir. The e/C128ect of boiling is relatively small, with most temperature decreasesbeing caused by dilution.
The reservoir fluid is significantly depleted in Mg and SO
4and enriched in K,
Ca, SiO 2and B compared with sea water. In the Fushime area, the sea water
heated at depth reacts with the host rock and then ascends by convection. Duringascent, dilution by marginal or shallow ground water occurs. During this cooling,K is precipitated, for example, as K-feldspar, and Ca is dissolved from CaSO
4
precipitated at shallower depths; anhydrite is observed at about 1000 m in depth.
The calculated Na/Cl ratios of the fluids are slightly lower than that of sea
water, but there is little change in the Na/Cl ratio during the dilution process. Onthe other hand, the K/Cl and Ca/Cl ratios are considerably larger than that of thesea water, and are variable. Fig. 11 shows cross plots of K/Cl and Ca/Cl ratiosagainst Clres. The K concentration of the fluid first increases by leaching of Kfrom the surrounding rock as the temperature increases. The K/Cl ratio thendecreases during ascent, i.e., the decrease in K concentration was larger than thatcaused by dilution. This depletion of K is likely caused by the precipitation of aK-bearing mineral such as K-feldspar or K-mica. Yagi (1990) notes that K-feldspar is abundant in the center of the feed points of the production wells.
On the other hand, the Ca/Cl ratio is also variable, but is more complicatedH. Okada et al. / Geothermics 29 (2000) 279–311 296
than the variation in the K/Cl ratio, and shows the opposite trend at lower Clres
value. This can be explained by the precipitation or dissolution of Ca-bearingminerals such as anhydrite or calcite. The Ca-bearing minerals show a retrograde
solubility, such that the Ca/Cl ratio will increase with temperature decrease. In the
case of the Fushime system, before sea water reached the hot and deep reservoir,calcium most likely precipitated as anhydrite. The calcium may then be leached bysubsequent dissolution of previously precipitated anhydrite or calcite during laterfluid ascent. However, the data do not lie on a single trend. Akaku and Reed(1995) suggested that a solid solution of epidote or other Ca-bearing minerals(e.g., anorthite) involved in dissolution may possibly reconcile the complexbehavior of Ca.
An Na–K–Mg triangular diagram (Fig. 12; Giggenbach, 1988) can represent the
extent and temperature of mineral equilibrium in the reservoir graphically, i.e., fullequilibrium, partial equilibrium, and immature solutions. All Fushime data fromwell discharges plot along the full equilibrium line; however, data of several hotsprings plot in the partial equilibrium or immature area. It is therefore thought
Fig. 10. Relationship between chloride concentrations and enthalpy calculated by the Na–K–Ca
geothermometer in the Fushime field. Broken line indicates the dilution by ground water. : Hot
springs, *: Geothermal fluids, P: Sea water, Clres: Reservoir Cl concentration recalculated by using
Na–K–Ca geothermometer.H. Okada et al. / Geothermics 29 (2000) 279–311 297
Fig. 11. Mole ratios of K 100/Cl and Ca100/Cl to the Clres (mg/kg) in the Fushime geothermal
fluids. Symbols are hot springs ( ), Fushime geothermal fluids ( *) and sea water ( P). Arrows show
geochemical processes in the reservoir.H. Okada et al. / Geothermics 29 (2000) 279–311 298
that the hot waters produced from the deeper part of the Fushime area reached
the full equilibrium condition.
Fluid that is produced from depths without the occurrence of alunite and
kaolinite should reach full equilibrium and have a near-neutral pH. However, in
the Fushime area, fluids discharged from several wells have a low pH despite thesuggestion of full equilibrium condition. However, there are no acid alterationminerals, such as alunite and kaolinite, in the core and cuttings of the Fushimefield. This discrepancy can be explained by precipitation of PbS, ZnS and CuSminerals inside the well and/or formation during discharge.
In comparison with other geothermal fields in Japan, the Fushime geothermal
fluid is rich in heavy metals, such as Pb and Zn. After a three-month discharge
test, some sulfide minerals such as galena or sphalerite were observed in the
pipelines or on the control valves, where rapid pressure and temperature dropsoccur. Akaku et al. (1991) showed that the following reaction occurs in a casingliner or a pipeline during fluid ascent:
MeCl
2H2S
MeS2Clÿ2H
Me:Pb,Zn,Cu
The solubility of sulfide minerals decreases with decreasing temperature, resulting
in the reaction shifting to the right; thus, sulfide mineral precipitation createsacidity.
Fig. 12. Na–K–Mg triangular diagram using geothermometer temperature recommended by
Giggenbach (1988). Symbols are hot springs ( ), the Fushime geothermal fluids ( *) and sea water
(P).H. Okada et al. / Geothermics 29 (2000) 279–311 299
The Fushime geothermal fluid most likely has a near neutral pH in the
reservoir. On the other hand, Reed (1991) computed the reservoir pH of the fluidby using the computer program CHILLER, and he suggested that the fluid had a
low pH due to contamination of small amounts of acidic gases such as HCl and
H
2SO4. Generally speaking, magmatic gases contain a small amount of HCl and/
or HF (Greenland et al., 1985). As discussed later, non-reactive gas compositions,such as N
2, He and Ar, show that the Fushime fluid contains very little magmatic
gas, the most likely source of HCl. Since no acidic alteration minerals such asalunite and kaolinite were found at the reservoir depth, we do not attribute theacidity of the Fushime geothermal discharge to magmatic gas. These observationsindicate that the fluids of the Fushime reservoir reached full equilibrium and that
acidic fluid does not exist in the reservoir.
4.2. Isotope compositions
ThedD and d
18O values of the fluids range from ÿ15.1 toÿ2.4-, and from
ÿ2.0 toÿ0.4-, respectively (Table 3). All fluid compositions plot to the right
(Fig. 13) of the local meteoric water line determined by Sakai and Matsubaya(1977). The d
18O value of local sea water is about 1 -lighter than that of the
SMOW value. The data also lie slightly to the right of the line connecting local
sea water and ground water. Some oxygen shift may have occurred in thereservoir due to rock–water interaction at high temperature.
In general, a thermal fluid that reacts with rock at high temperature becomes
enriched with oxygen isotope (
18O) (Fritz and Fontes, 1980). However, the oxygen
shift is relatively small in this field. There are two possibilities to explain thisphenomenon; the reaction time between rock and sea water is relatively short, orthe rock/water ratio is smaller than that of ordinary geothermal fields.
4.3. Gas chemistry
The Fushime fluids contain noncondensible gases with a concentration less than
1.0 vol% in steam. The dominant gas is CO
2, ranging from 17.3 to 284 mmol/kg
in steam, and next highest concentration is that of hydrogen sulfide (Table 3).Nitrogen, hydrogen and methane occur in small amounts. It is very important fora reservoir evaluation to estimate steam fraction in the reservoir. Giggenbach
(1980) demonstrated that by using equilibria for the same gas reaction, it was
possible to evaluate graphically both steam fraction and reservoir temperature inthe reservoir. To estimate the steam fraction in the reservoir, the data were plottedon the Giggenbach diagram (Fig. 14). All wells except for Nos. 11 and 12 wellshave excess steam. Most data, except for that from one well, plot in a region inwhich the steam fraction ranges between several percent and about 10%. Since theexception, well 10, has a much higher CO
2ratio and gas volume, the data of this
well plot at a steam fraction of as much as 50% or more. However, it is not
thought that this value shows the correct steam fraction because the total enthalpyH. Okada et al. / Geothermics 29 (2000) 279–311 300
is not much higher than that of other wells. Thus, the gas analysis of this well
may be in error.
We can generally deduce a gas origin by plotting relative N 2, He and Ar
contents in a triangular diagram: namely, meteoric, magmatic and crustal. Fig. 15
shows the triangular diagram on which these gas data are plotted, where all threedata fall close to the point of air or air-saturated meteoric water. Giggenbach(1992) demonstrated that data plotting near air indicate air contamination duringor after sampling. Since we eliminate other samples including oxygen to plot onthe graph, air contamination to these samples during or after sampling is unlikely.The compositions of the non-reactive gases, N
2, Ar, and He, therefore suggest
that very little crustal or magmatic gas is present in the fluids of this area, but
that these non-reactive gases are dominated by air-saturated ground water.
Fig. 13. Relationship between dD and d18O in the Fushime geothermal fluids (relative to SMOW). *:
Fushime geothermal fluids, P: Local sea water. Broken line represents the mixing line between local
sea water and local meteoric water. The solid line represents the local meteoric line reported by Sakai
and Matsubaya (1974). The shadow area indicates the ground water composition in the Fushime field.H. Okada et al. / Geothermics 29 (2000) 279–311 301
5. Geothermal structure
5.1. Temperature distribution
The standard series of temperature loggings in this field consists of 42-h
temperature recovery tests before and after cooling by water injection, and
Fig. 14. Variation of the methane and ammonia concentration quotients, K0c, in relation to the gain or
loss of steam underground; y= steam fraction (from Giggenbach, 1980). K0c= XCO 2XH 2/XCH 4.
This equation is based on the following chemical reaction: CH 4+2H 2O=CO 2+4H 2. The data collected
in 1990 are plotted on the graph.H. Okada et al. / Geothermics 29 (2000) 279–311 302
additional temperature surveys conducted afterwards, e.g., before flow tests.
Typical temperature profiles from several wells of the Fushime field and thesurrounding area are shown in Fig. 16. There is a temperature inversion at about
1 km depth inside the structural depression. The temperature of the inversion zone
decreases to about 100–150 8C, with the inversion clearly indicated in the high
temperature area. Below the inversion zone, the temperature reaches about 350 8C
at the production depth. The shallow high-temperature zone above the inversionzone is about 200–250 8C. Outside the structural depression, the temperature
profiles show quite a di/C128erent pattern from those inside, with a small thermalgradient, lower temperature and no inversion.
Comparing this temperature distribution to the geologic structure (Fig. 6), it
appears that hot fluid ascends from the area around the dacite intrusion, located
in the most depressed part, and that convection is focused in relatively porous tu/C128and tu/C128 breccia of the Yamagawa formation. On the other hand, outside thestructural depression there is little ascent of hot fluid. The trend of increasingtemperature towards Takeyama supports the idea that the heat source for thisfield lies below Takeyama volcano (Yoshimura et al., 1988), which has a fission-track age of 26,000 year B.P. (Kamitani et al., 1976).
Fig. 15. Relative N 2, He and Ar contents in geothermal gas discharges on molar basis. Sampling date is
di/C128erent from that given in Table 3. Numbers shown indicate well number.H. Okada et al. / Geothermics 29 (2000) 279–311 303
Fig. 16. Temperature profiles from wells drilled in the Satsunan area. Temperature profiles from the
Fushime field show quite di/C128erent pattern from wells in the surrounding area.H. Okada et al. / Geothermics 29 (2000) 279–311 304
5.2. Permeability distribution
Fluid loss during drilling is one of the most reliable indicators of permeability
in the well. Two major zones of considerable loss were recognized during drilling
(Fig. 17). One is a NNE–SSW trending loss zones to the West, and the other is aNW–SE trending loss zone to the South. The western loss zone is reflected in anelectrical discontinuity detected by the Tubel resistivity survey. The southern losszone is concordant with the TTL. The distribution of these two loss zonesindicates that they are vertical in orientation, suggesting they are related to faults.In Fig. 7, the relationship between loss rates and the TTL is shown. It is clear thatlarge losses tend to occur near the TTL.
To examine the features of the TTL, we carried out an FMI
1(Fullbore
Formation MicroImager) log in addition to standard logging in well 23. FMIimage data were acquired over the completion interval of the well, between 811and 1189 m depth. The image was used in conjunction with drill cuttings andother logging data to determine the depth of lithologic changes, and to identifyfractures across the borehole. Most of the FMI images acquired indicated tu/C128breccia, which was easily recognized. Andesite dikes appeared as resistive and highvelocity zones and showed a less clear texture in the FMI image. More than 250
fractures/boundary features were picked up in the FMI analysis, and these were
classified into five categories; drilling-induced fractures, open fractures, bedboundaries, healed fractures, and poor quality (uncertain) fractures. Although thedistribution of dikes occurring in wells 7, 15 and 23 indicates a clear NWdistribution trend, it was rather dicult to identify the direction of each dike thatappeared in the FMI image. Most natural fractures, i.e. open fractures plus healedfractures, show a NW–SE direction (Fig. 18), and their density (number offractures over 10 m interval) clearly increases in the dike zone. Drilling-induced
fractures also display a NW–SE strike direction in the interval of dike occurrence.
These observations indicate that the fractures formed as a result of dike intrusionalong the TTL. These permeable zones are presently utilized as the mainreinjection zones for this field.
The permeability for the production zone of the field is more complicated.
When drilling into the high-temperature reservoir, we changed drilling fluid frommud to water to avoid formation damage (Yoshimura and Ito, 1994). In mostcases, circulation loss occurred and the rate gradually increased after the start of
water drilling, which made it rather dicult to identify the depths of major losses
that could become feed zones. Flowmeter logging and temperature surveys beforeand after cold water injection helped to indicate the permeable zones in thecompleted section of the production wells.
The major feed zones in this field usually occur around the dacite intrusion.
Determining the distribution of the dacite intrusion is therefore the key toexploration in this field. The information to date indicates that the intrusive body
1Trade mark of Schlumberger.H. Okada et al. / Geothermics 29 (2000) 279–311 305
Fig. 17. Production and reinjection areas in the Fushime field. Main production area is defined as the
area of >300 8C at a depth of 1700 m below sea level; the reinjection area is defined as the crossing
area of the Takeyama–Tsujinodake line (TTL) and NE trending fault zone detected by Tubel resistivitysurvey.H. Okada et al. / Geothermics 29 (2000) 279–311 306
is located in the central area of the Fushime depression. The tu/C128s surrounding this
dacite intrusive body have commonly undergone severe silicification and containpyrite within about 40–100 m of the contact, indicating the presence of
hydrothermal fluid. An example of a log from a production well is shown in Fig.
19. X-ray analysis of drill cuttings from this zone show a clear increase in thecontent of K-feldspar and decrease in that of plagioclase (Yagi, 1990).Concordantly, the gamma-ray log shows a high anomaly in this feed zone.
6. Conclusions
Geologic analysis showed that the main reservoir for the Yamagawa power
station in the Fushime geothermal field occurs in a fractured zone developedaround a dacite intrusion. A knowledge of the distribution of the dacite intrusionis therefore the key to exploration in this field. The information to date indicatesthat the intrusive body is located in the central area of the Fushime depression.
Fig. 18. Fullbore Formation MicroImager (FMI) analysis in the dike zone. Data are acquired from
well 23. Rose diagram shows the strike of open fractures and healed fractures. Fracture density is thenumber of each fracture picked over a 10 m interval, which shows clear increase in the dike zone.Diagram of poles to planes is plotted on upper hemisphere of Wul/C128 stereonet.H. Okada et al. / Geothermics 29 (2000) 279–311 307
Fig. 19. Example of a log from a typical production well. The main feed zone is located near the top of the dacite.H. Okada et al. / Geothermics 29 (2000) 279–311 308
We defined the main production zone for the power station as the area where
temperatures exceed 300 8C at a depth of 1700 m below sea level (Fig. 17). In some
wells a high temperature of over 350 8C has been measured at the production
depth.
Chloride concentration of local sea water is 19,800 mg/l. Based on geochemical
information, the reservoir chloride concentration of this field ranges from 11,600to 22,000 mg/kg. The Clres, Br/Cl ratios and stable isotope data indicate that theFushime geothermal fluid originated from heated sea water and is diluted byground water during ascent. Some fluids produced from geothermal wells have alow pH of about 4. It is thought that sulfide mineral precipitation (PbS, ZnS)during production causes this acidic fluid.
High permeability zones identified in serious lost-circulation zones are
associated with faults. One of these is clearly related to the NW–SE trending TTLassociated with an andesite dike swarm. Another fault trends NNE–SSW and isreflected by electrical discontinuity. The reinjection area is defined as theintersection of these structures at a distance from the production zone (Fig. 17).
The completion intervals for the production wells range from about 1400 to
2100 m below sea level, whereas the depth of the reinjection zone is betweenabout 800 and 1200 m depth. This also helps to keep the production and
reinjection zones separate.
Acknowledgements
The authors would like to thank Dr M. Sasada of the Geological Survey of
Japan for providing the opportunity to submit this paper, and the JapanPetroleum Exploration Co. (JAPEX), Ltd and JAPEX Geothermal Kyushu(JGK), Ltd for their permission to publish the paper. The authors also thank MrK. Akaku and Mr O. Nakagome of JAPEX and Mr K. Sakuma of JGK forhelpful discussions.
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Okada (2000) geology and fluid chemistry of fushime geothermal.txt
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Metamorphic evolution of kyanite–staurolite-bearing
epidote–amphibolite from the Early Palaeozoic Oeyama belt,
SW Japan
T. TSUJIMORI1,2AND J. G. LIOU2
1Research Institute of Natural Sciences, Okayama University of Science, Okayama 700–0005, Japan
(tatsukix@pangea.stanford.edu)
2Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305, USA
ABSTRACTEarly Palaeozoic kyanite–staurolite-bearing epidote–amphibolites including foliated epidote–amphi-
bolite(FEA),andnonfoliatedleucocraticormelanocraticmetagabbros(LMG,MMG),occurintheFukoPassmetacumulateunit(FPM)oftheOeyamabelt,SWJapan.Microtexturalrelationshipsandmineralchemistrydefinethreemetamorphicstages:relictgranulitefaciesmetamorphism(M1),high- P
(HP)epidote–amphibolitefaciesmetamorphism(M2),andretrogression(M3).M1ispreservedasrelict
Al-richdiopside(upto8.5wt.%Al
2O3)andpseudomorphsafterspinelandplagioclaseintheMMG,
suggestingamedium- Pgranulitefaciescondition(0.8–1.3GPaat>850 /C176C).Anunusuallylow-variance
M2assemblage,Hbl+Czo+Ky±St+Pg+Rt±Ab±Crn,occursinthematrixofallrock
types.ThepresenceofrelictplagioclaseinclusionsinM2kyaniteassociatedwithclinozoisiteindicatesahydrationreactiontoformthekyanite-bearingM2assemblageduringcooling.Thecorundum-bearing
phaseequilibriaconstrainaqualitativemetamorphic P–Tconditionof1.1–1.9GPaat550–800 /C176Cfor
M2.TheM2mineralswerelocallyreplacedbyM3margarite,paragonite,plagioclaseand ⁄orchlorite.
ThebreakdownofM2kyanitetoproducetheM3assemblageat<0.5GPaand450–500 /C176Csuggestsa
greenschist facies overprint during decompression. The P–Tevolution of the FPMmay represent
subductionofanoceanicplateauwithagranulitefacieslowercrustandsubsequentexhumationinaPacific-typeorogen.
Key words:EarlyPalaeozoic;epidote–amphibolite;kyanite;polyphasemetamorphism; P–Tpath;
staurolite.
INTRODUCTION
Amphibolites are common in many Barrovian-type
metamorphicterranes.Mostcontainhigher-variance
mineral assemblages that do not constrain wellthe metamorphic conditions. However, some high-
pressure(HP)amphiboliteswithanaluminousbulk-
rockcompositioncontainkyaniteandstaurolite(e.g.Gibson,1979;Selverstone etal.,1984;Yokoyama&
Goto, 1987; Gil Ibarguchi etal., 1991; Kuyumjian,
1998). Such unusual amphibolite assemblages areusefultoconstrainthe P–Tpath(e.g.Arnold etal.,
2000)asreactiontexturesarewelldisplayedinalumi-
nous mineralssuchaskyanite(e.g.Spear&Franz,1986;Cotkin etal.,1988).
Early Palaeozoic kyanite- and staurolite-bearing
epidote–amphibolite and metagabbro of the Fuko
Pass metacumulate unit (FPM) occur in the peri-dotite body of Early Palaeozoic Oeyama belt
(Kurokawa, 1975; Kuroda etal., 1976; Tsujimori,
1999).TheKy–St-bearingrockshavebeenconsideredto have recrystallized at medium- P(c.0.5 GPa),
possibly in an ocean-floor setting, as a cumulatemember of an ophiolitic succession (Kurokawa,
1985).Ontheotherhand,theHbl+Czo+Ky+Pg+Rt assemblage suggests HP metamorphism(Tsujimori, 1999). Moreover, medium- Pgranulite-
faciesrelicsofAl-richdiopsideandpseudomorphsof
spinel and plagioclase were recently identified insome metacumulate rocks (Tsujimori & Ishiwatari,
2002).Theserocksconstrainauniquemetamorphic
evolution from the granulite-facies to HP meta-morphism. The occurrence of an Early Palaeozoic
HPmetamorphiceventintheFPMmaysignifythe
earliestsubduction-metamorphisminthePacific-typeorogenoftheJapaneseIslands.
Thispaperpresentsnewpetrologicaldataforthe
Ky–St-bearing epidote–amphibolites and metagab-
bros of the FPM, which are coupled with previousdatatoevaluateapolyphasemetamorphicevolution
oftheFPManddiscussthetectonicsignificanceof
thisevolution.MineralabbreviationsareafterKretz(1983) and the term /C212hornblende (Hbl) /C213is used to
describe Ca-amphibole with dominantly pargasitic,
tschermakitic and edenitic composition throughoutthispaper.J. metamorphic Geol. , 2004, 22,301–313 doi:10.1111/j.1525-1314.2004.00515.x
/C2112004BlackwellPublishingLtd 301
GEOLOGICAL SETTING
The Oeyama belt is an Early Palaeozoic ophiolitic nappe that
occupiesthehigheststructuralpositioninthePhanerozoicPacific-
typeorogenofsouth-westJapan(e.g.Ishiwatari&Tsujimori,2003)(Fig.1). Several serpentinized peridotite bodies in the ChugokuMountainsconsistmainlyofresidualperidotitewithminorgabbrodykes and podiform chromitites (Arai, 1980; Kurokawa, 1985;Matsumotoetal.,1997).Theperidotitehasalherzoliticcomposition
in the eastern part, but is harzburgitic in the western part. The
gabbrodykeshaveMORB-likeaffinityandyieldSm ⁄Ndisochron
agesofc.560Ma(Hayasaka etal.,1995),suggestingaCambrianor
earlieragefortheophioliteformation.TheOeyamaperidotitemayhavebeenderivedfromsuprasubductionzonemantlebeneathanintraoceanicarc(Tsujimori&Itaya,1999;Ishiwatari&Tsujimori,
2003).
TwodifferentHPmaficrocksareassociatedwithserpentinized
peridotitebodiesintheOeyamabelt.TheyoungerHProcksconsistof330–280Mablueschisttoeclogitefaciesmetasedimentsandminormetabasalt,andrepresentfragmentsoftheLatePalaeozoicRenge
blueschist nappe underlying the Oeyama belt (Nishimura, 1998;Tsujimori,1998;Tsujimori&Itaya,1999).Incontrast,theolderHProcksarecharacterizedby470–400Madeformedmetagabbroandmetaclinopyroxenite(Nishimura&Shibata,1989;Tsujimori etal.,
2000).TheFPMbelongstothelattertype.
TheFPMoccursasafault-boundedsliceatthetopographically
higher portion of the Oeyama peridotite body that tectonicallyoverlies the Permian accretionary complex of the Akiyoshi belt(Fig.1).TheFPMisameta-ultramafic–maficcomplexconsistingofmetamorphosedclinopyroxenite,wehlriteandcumulategabbro(Kurokawa, 1985). On the other hand, the Oeyama peridotitebody is composed of fertile harzburgite with minor dolerite and
gabbro dykes (Uda, 1984; Kurokawa, 1985). Both have been
highlyserpentinized,andsubsequentlysignificantlyoverprintedbycontactmetamorphismaroundLateCretaceousgraniticintrusions.Uda(1984)identifiedfivecontactmetamorphiczones:I–Atg;II–Ol+Atg+Di; III – Ol+Atg+Tr; IV –Ol+Tlc±Tr±Cum;andV–Ol+En±Tr±Hbl.Primaryigneousphases,OlFig. 1.(a)AsimplifiedmapoftheChugokuMountains,showingvariouspre-Triassicpetrotectonicunitsandthelocalityofthe
Oeyamaarea.(b)GeologicalmapoftheOeyamaarea,showingsamplelocalityoftheKy–St-bearingrocks(afterTsujimori etal.,
2000).302 T. TSUJIMORI & J. G. LIOU
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(Fo90)91)+Opx (Al 2O3¼3.2–4.3 wt.%)+Cpx (Na 2O<0.3
wt.%)+Cr-Spl (Cr# ¼0.34–0.38), are well preserved in Zone I
rocks(Kurokawa,1985;Tsujimori,1999).TheFPMextendsfrom
Zone II to Zone V; few outcrops lack the thermal effect of thecontact metamorphism. No gabbro intrusions cut through theFPM, whereas the adjacentserpentinized peridotite was intrudedbysyn-orpost-tectonicgabbroicrocks.Thus,itisinterpretedthattheFPMwastectonicallyjuxtaposedwiththeOeyamaperidotite
body. K ⁄Ar hornblende ages of the FPMyield 443–403M a
(Tsujimorietal.,2000).
Sampleswerecollectedatanoutcrop(20 ·3m)neartheFuko
Pass where the effect of contact metamorphism is minimal (Ku-rokawa,1985).Ky–St-bearingrocksoccurtogetherwithirregularblocksofmassivecoarse-grainedmetagabbro,virtuallyundeformed,inthematrixoffoliated(orgneissose)epidote–amphibolite.Tsuji-
mori(1999)usedtheterm /C212metagabbro /C213toindicateanyundeformed
epidote–amphiboliteinwhichoriginalgabbroictexturesarewell-preserved,anddistinguisheditfromthe /C212foliatedepidote–amphibo-
lite /C213.Themetagabbroissubdividedintoleucocraticandmelanocratic
varietiesbythecontentofclinozoisite.
ANALYTICAL METHODS
Concentrationsofmajor(Si,Ti,Al,Fe,Mn,Mg,Ca,Na,K&P)andtrace(Ni,Cu,Zn,Pb,Y&V)elementswereanalyzedbyaRigakuSystem3270X-rayfluorescencespectrometerwithRhtubeatKanazawaUniversity.Theoperatingconditionsforbothmajorandtraceelementswere50kVacceleratingvoltageand20mAbeamcurrent.Othertraceelements(Sc,Cr,Co,La,Sm,Eu,Lu&Th)were
determined by instrumental neutron activation analysis (INAA
method).TheINAAsampleswereactivatedatKyotoUniversityReactor,andthegamma-rayspectroscopicanalysesweredoneattheRadioisotopeLaboratoryofKanazawaUniversity.
Electron microprobe analysis was carried out with a JEOL
JXA-8800R at Kanazawa University and JEOL JXA-8900R atOkayamaUniversityofScience.Thequantitativeanalysesofrock-
formingmineralswereperformedwith15kVacceleratingvoltage,
12nAbeamcurrentand3–5 lmbeamsize.Naturalandsynthetic
silicates and oxides were used for calibration. The ZAF method(oxidebasis)wasemployedformatrixcorrections.
PETROGRAPHY
Threerocktypesaredistinguishedaccordingtotheir
lithological and petrographical features: (1) foliatedepidote–amphibolite(FEA);(2)leucocraticmetagabbro(LMG); and (3) melanocratic metagabbro (MMG).
Althoughepidote–amphibolitefaciesassemblagesare
dominantinalllithologies,eachrocktypepreservestextualevidenceformineralgrowthduringanearlier
stage and of retrograde metamorphic events. Three
metamorphicstages,M1,M2andM3,aredistinguishedonthebasisofmicrotexturalrelationshipsandmineral
chemistry.Mineralparagenesesfordifferentmetamor-
phicstagesoftheserocktypesaresummarizedinFig.2.M2definesthepeakHPmetamorphismofallrocktypes,
whereasM1mineralsareidentifiedasrelictminerals,
particularlygranulitefaciesrelicsandpseudomorphsin
melanocraticmetagabbro. M3representslower- Pret-
rogressionduringdecompression.
FEA: foliated epidote–amphibolite
Thisrocktypeconsistsmainlyofhornblende,clino-zoisiteandkyanitewithsmallamountsofparagonite,rutile,chlorite,margarite,staurolite,corundum,zoi-
site,andrarealbiteandmuscovite.Ilmenite,sulphides
andapatiteoccurasaccessories.Afoliationdefinedbypreferred orientation of nematoblastic hornblende,
kyanite and paragonite is developed. Mosaic aggre-
gatesofclinozoisiteareintergrownwithparagonite,kyanite,rutile,hornblendeandstaurolite;thesemin-
eralsalsooccurasinclusionsinclinozoisite(Fig.3a,b).
Rare zoisite blebs (<0.1mm), albite and musco-vite are included in clinozoisite. Some staurolite
(<0.5mm)occursasaggregates(upto5mm)with
minorcorundum(<0.1mm)andraremagnetite;theaggregateisarmouredbykyaniteandclinozoisiteatthe margin (Fig.3c). These matrix minerals and
inclusionsinclinozoisitearethoughttobeintextural
equilibrium,andthemineralassemblageHbl+Czo+Ky±St+Pg+Rt±Ab±Crncharacterizesthe
peakM2.Somekyanitecrystalscontainplagioclaseas
tinyinclusions(<0.03mm)(Fig.3d);thismayrep-resentaprecursorofthepeakofM2,possiblyarem-
nantoftheM1stage.RutilecontainsexsolutionblebsGS EA GRFEA & LMG MMGHORNBLENDEMINERALFACIES
KYANITE
CORUNDUM STAUROLITE
ALBITE PLAGIOCLASE
P
SPARAGONITE
RUTILESTAGE M 2 M1 M3
CHLORITEMUSCOVITEMARGARITE
ILMENITEZOISTECLINOZOISITE
EPIDOTE
HORNBLENDE
KYANITE(P)
CORUNDUM(S)MAGNETITE(S)
PARAGONITE(P)CLINOPYROXENE
MARGARITE
ILMENITE
PSEUDOMORPHS AFTER PLAGIOCLASE (P) AND SPINEL (S) MAJOR PHASESRUTILECHLORITEZOISTE(P)CLINOZOISITE(P)TREMORITE
EPIDOTE
PSEUDOMORPHSMINOR PHASES
Fig. 2.Mineralparagenesesforthreestagesofmetamorphic
recrystallizationforfoliatedepidote–amphiboliteandleucocraticmetagabbro(FEA&LMG),andmelanocraticmetagabbro
(MMG).KYANITE–STAUROLITE-BEARING EPIDOTE–AMPHIBOLITE 303
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ofilmenite;somearerimmedbythinilmenite.Retro-
gradeM3whitemicaoccursascoroniticaggregates
around M2 kyanite (Fig.3d). Chlorite and tiny
plagioclase(<0.05mm)occurrarelyasM3mineralswithM3margaritearoundM2kyanite.M3chlorite
replaces M2 hornblende along the cleavage and
stauroliteinaggregates.TheassemblageMrg+Pg±Ms+Chl±Pl characterizes the M3 retrogression.
Insomesamples,epidoteovergrowthsoccuratinter-
faces between M2 clinozoisite and M3 hydrousminerals.EpidotemaybeaproductofM3orlater.
LMG: leucocratic metagabbro
This rock type consists mainly of clinozoisite and
kyanitewithminorhornblende,margarite,paragonite,
staurolite,rutile,chlorite,zoisiteandmuscovite.Rareilmeniteandapatiteoccurasaccessories.Retrogression
inthisrockismoreobviousthanintheFEA.Coarse-
grained clinozoisite (up to 15mm) contains all M2mineralsasinclusions.Somekyanitegrainsshowpale-
bluepleochroism.StauroliteincludestraceM2parag-
onite.KyaniteispseudomorphedbyM3margariteandparagonite(Fig.3e,f).M3muscoviteoccurswithM3margarite and paragonite (Fig.4). M3 margarite is
associatedwithM3chloriteandrareM3plagioclase.
MMG: melanocratic metagabbro
Thisrocktypeconsistsmainlyofhornblende,clino-
pyroxene, corundum-magnetite symplectite, clinozoi-
siteandchloritewithminorkyanite,margarite,rutile
andparagonite.ClinopyroxeneisarelictmineralofM1;itoccursasequigranulargrains(3–6mm)partly
replacedbyM2hornblende(Fig.3g).TheCrn–Mag
symplectitehasequidimensionalshape(1–3mm),andconsistsmainlyofintergrowthsofcorundumandgra-phicmagnetitewithminorZn-richspinel(Fig.3g,h);it
wasinterpretedtobeapseudomorphafterspinelof
the M1 assemblage (Tsujimori & Ishiwatari, 2002).Moreover,mosaicaggregatesofclinozoisiteincluding
kyanite and paragonite at interstices between M1
clinopyroxene and Crn–Mag symplectite were inter-pretedtobeapseudomorphafterplagioclaseoftheM1assemblage.Mostkyaniteinclusionsinclinozoisiteare
replacedbyM3margarite.M3chloritereplacesM2hornblende,andoccursaschlorite-richclots(3–5mm),
or aggregates surrounding the corundum-magnetite
symplectite.PrismaticM3epidote(<1mminlength)and tiny magnetite are randomly oriented and fre-
quentlyassociatedwiththechloriteclots.
Bulk-rock chemistry
OneFEA,twoLMGandtwoMMGsampleswereselectedforgeo-
chemicalstudy.AnalysedbulkcompositionsarelistedinTable1.MajorelementanalysesshowthatalllithologiescontainextremelylowSiO
2(38.7–41.1wt.%),highAl 2O3(18.9–25.5wt.%),CaO(13.1–
15.3wt.%),andlowNa 2O+K2O(<2.2wt.%)with48–66mol.%
normativeanorthiteand16–28mol.%olivine.TheMg ⁄(Mg+Fe*)
ratios range from 0.45 to 0.65 and increase in the order ofFEA<LMG<MMG;Fe* ¼totalFeasFe2+.TheFEAand
LMGareenrichedinSr(402–479p.p.m.)whereastheMMGshowsenrichmentsofNi(97–99p.p.m.),Sc(135–143p.p.m.),andZn(166–224p.p.m.).TheN-MORBnormalizedtraceelementabundancesof
theanalysedsamplesagainstthestandardvaluesofSun&McDo-
nough(1989)areshowninFig.4forcomparison.Theyarecharac-terizedbyahighconcentrationoflargeionlithophile(LIL)elements,andshowsimilarpatternstothosefrommeta-anorthositeoftheWesternGneissRegion,Norway(Cotkin,1997)andkyanite–staur-olite-bearingrockfromtheCaboOrtegalcomplex,NWSpain(GilIbarguchietal.,1991).ThelowREEcontentsoftheanalysedsamples
suggestacumulativeorigin.Thebulkrockcompositionandthecal-
culatednormsuggestanoriginfromatroctolitic-andanorthositiccumulatewithabundantanorthite-richplagioclase.
MINERAL CHEMISTRY
Representative electron microprobe analysis of rock-forming mi-neralsintheFPMarepresentedinTable2.
Ca-amphibole
ThestructuralformulaeofamphibolearecalculatedbasedonO ¼23
and the Fe2+⁄Fe3+ratio was estimated on the basis of total
cation ¼13,excludingCa,NaandK(Leake etal.,1997).Most10-210-1101
0102103
Rb Ba Th K Ta Nb La Ce Sr P Zr Hf Sm Ti Y Yb LuMeta-anorthosite, Norway (Cotkin, 1997)
Ky-St ultrabasic rock, Spain (Gil Ivarguchi, 1991)FEA
LMGMMG
Fig. 4.N-MORBnormalizedincompatibleelementpatternsof
analyzedsamplesofFEA,LMGandMMG.Forcomparison,theelementabundancepatternsofmeta-anorthositeoftheWesternGneissRegion(Cotkin,1997)andkyanite–staurolite-bearingrockfromtheCaboOrtegalcomplex(NWSpain)(Gil
Ibarguchietal.,1991)arealsoillustrated.
Fig. 3.PhotomicrographshowingmicrotextureoftheKy–St
bearingrocks.(a)TheoccurrenceofM2kyaniteoftheFEA.(b)M2paragoniteinclusionsinclinozoisiteoftheFEA.(c)Stauroliteaggregateassociatedwithkyanite,clinozoisiteandhornblendeoftheFEA;stauroliteisassociatedwithtinycorundum.(d)PlagioclaseinclusionsinM2kyaniteinequilib-
riumwithHblandCzooftheFEA;M3assemblageMrg+Pg
developedaroundKy.(e)Coarse-grainedpseudomorphcon-sistingofM3assemblageofMrg+PgafterM2kyaniteoftheLMG.(f)M3assemblageChl+Mrg+PgdevelopedaroundM2Ky+St.(g)RelictM1clinopyroxeneandpseudomorphofCrn+MagafterspineloftheMMG;M2hornblende
replacesrelictM1CpxandM2clinozoisiteoccursbetweenCpx
andtheCrn-Magsymplectite.(h)Enlargedviewofcorundum-magnetitesymplectiteafterM1spineloftheMMG.KYANITE–STAUROLITE-BEARING EPIDOTE–AMPHIBOLITE 305
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/C212hornblende /C213of the studied rocks are tschermakitic to pargasitic
compositionscontainingupto18wt.%Al 2O3,3.4wt.%Na 2Oand1.2
wt.%TiO2,whereasactinoliticortremoliticcompositionsareidenti-
fiedat therim of relativelycoarse-grainedcrystals intheMMG(Fig.5).IndividualhornblendegrainsintheFEAandLMGareun-zoned,buttheircompositionsarevariableatthethinsectionscale.
MinimumX
Mg[¼Mg⁄(Mg+Fe2+)]valuesbecomehigherinthe
orderofFEA(0.59)<LMG(0.68)<MMG(0.73).HornblendeoftheMMGcontainssignificantlylowerM4-site([M4]Na)thanthatinthe
FFMandLMG,apparentlyduetodifferentbulkrockcomposition.
Epidote group minerals
TheXFe3+[¼Fe3+⁄(Fe3++Al)]ofM2clinozoisiteinalllithologies
variesfrom0.10to0.24.The XFe3+ofazoisiteblebis0.03–0.07.
M3epidoteoftheFEAandMMGischaracterizedbydistinctlyhighX
Fe3+(0.26–0.29).
Kyanite
Pale-bluecolouredkyaniteintheLMGcontainsupto1.3wt.%Fe
2O3,whereasothercolourlesskyanitehas<0.8wt.%Fe 2O3.
White mica
ChemicalcompositionsofwhitemicaareshowninNa(Pg)–K(Ms)–Ca(Mrg)diagramsofFig.6.M2paragoniteoftheFEAcontainsslightly higher Mrg component (Mrg ¼11–18mol%) than that
of the LMG and MMG (Mrg ¼4–6mol%); Ms component is
<10mol%.M2paragoniteinclusionsinstauroliteoftheLMGare
characterized by extremely high Mrg component (Mrg ¼
36–47mol%).M2muscoviteoftheFEAischaracterizedby6.2–6.3Sip.f.u.(O ¼22)and17–24mol%Pg.M3paragoniteoftheFEA
andLMGcontains2–14mol%Mrg.M3margariteshowsawidecompositional range; Pg content varies from 7 to 34mol%.
M3 muscovite is characterized by 6.0–6.2 Si p.f.u. (O ¼22) and
<10mol%Pg.
Staurolite
The structural formulae of staurolite were calculated based onO¼46.StauroliteoftheFEAcontainsaslightlyhigher X
Mg(0.19–
0.36)butlowerZnO(0.7–1.6wt.%)thanthatofLMG( XMg¼0.16–
0.29, ZnO ¼0.9–2.0 wt.%). The analysed staurolite shows good
negativecorrelationof(Fe+Mg+Zn)against(Si+Ti+Al>8) (Fig.7). The correlation overlaps a substitution line of(Mg+Fe)
3+1.5xAl18-xSi8O46that represents magnesian staurolite
(upto0.61XMg)incorundum-bearingamphibolitefromaneclogitic
metagabbrounitoftheSambagawabelt,SWJapan(Yokoyama&Goto,1987).
Feldspar
M1 plagioclase inclusion in M2 kyanite of the FEA contains26–34mol% An. M2 albite in clinozoisite of the EFA contains<5mol%An(maximum12mol%).M3plagioclasevariesfrom18to38mol%An.
Clinopyroxene
M1clinopyroxeneoftheMMGisAl-richdiopsidethatcontainsupto8.5wt.%Al
2O3,upto0.9wt.%TiO 2and<0.34wt.%Na 2O.
TheXMgvariesfrom0.78to0.94.TheAlcontentdecreasesgradually
towardthecontactwithclinozoisite.
Chlorite
M3chloriteoftheFEAandLMGischaracterizedbylowSi(5.3–5.5p.f.u.forO ¼28)andhighAl(5.1–5.8p.f.u); X
Mgis0.71–0.77inthe
FEA,and0.64–0.77intheLMG.IntheMMG,M3chloritearoundthe corundum–magnetite symplectite is aluminous (Al p.f.u. ¼
5.6–5.9,O ¼28)withX
Mgrangingfrom0.70to0.75,whereasthat
in chlorite-rich clots shows a wide compositional range (Al
p.f.u. ¼5.0–5.9,XMg¼0.65–0.87).
Others
Rutile contains 0.3–0.8 wt.% Fe 2O3. Corundum in staurolite
aggregatesoftheFEAcontains0.8–1.2wt.%Fe 2O3.Corundumin
thecorundum-magnetitesymplectiteoftheMMGcontains0.6–1.7wt.%Fe
2O3and<0.1wt.%Cr 2O3.Magnetiteintergrownwiththe
corundumcontains0.2–1.4wt.%Al 2O3and0.2–0.4wt.%TiO 2.Zn-
bearingspinel commonlyshows exsolutiontexture;Ti-rich spinellamellae(0.5–6.0wt.%TiO
2,7–14wt.%ZnOand XMg¼0.24–0.51)
aredevelopedparalleltothe{100}planeofTi-poorspinel(<0.2wt.%TiO
2,10–21wt.%ZnOand XMg¼0.43–0.92).Ilmeniteinthe
corundum-magnetitesymplectitecontains2–7wt.%MnO,whereas
thatintheFEAandLMGcontains<2wt.%MnO.
METAMORPHIC CONDITIONS
Basedontheobservedpetrographicfeatures,acounter
clockwiseP–Tevolutionthatpassesfromamedium- P
granulitefaciesstage(M1)throughaHPmetamorphicRock type FEA
Major-element compositions (in wt %)
SiO 2 41.08 39.51 39.37 38.86 39.04
TiO 2 0.78 0.50 0.65 0.56 0.52
Al2O320.63 25.46 23.97 19.41 18.92
Fe 2O3* 12.78 10.44 11.02 12.31 11.75
MnO 0.16 0.11 0.12 0.10 0.10
MgO 6.74 4.38 5.03 11.37 11.19CaO 14.08 15.29 15.22 13.09 14.28
Na
2O 1.84 0.91 1.00 0.88 0.82
K2O 0.36 0.58 0.47 0.14 0.11
P2O5 0.05 0.01 0.01 0.01 0.01
Total 98.45 97.18 96.85 96.72 96.73
Trace-element compositions (in ppm)Sc 43.3 30.1 34.5 143.1 135.5
V 536 415 425 633 610
Cr 33 59 76 64 52
Co 61 48 40 74 78
Ni 18 14 18 97 99
Cu 123 273 252 109 105
Zn 92 62 71 224 166
Rb 6 16 10 n.d. n.d.
Sr 402 479 411 147 139
Y 20 10 12 12 11
Zr 25 15 15 21 22
Ba 156 192 157 144 117
La 2 n.d. 1 1 1
Sm 21 11 1
Eu 0.9 0.5 0.6 0.5 0.6
Lu 0.3 0.1 n.d. 0.2 n.d.
Th 1.0 3.0 n.d. 4.0 n.d.
*Total Fe as Fe
2O3.
n.d. = not detected.MMG LMGTable 1.Bulk-rockcompositionsoftheFPM.306 T. TSUJIMORI & J. G. LIOU
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Table 2.Representativeelectron-microprobeanalysesofrock-formingmineralsintheFPM.KYANITE–STAUROLITE-BEARING EPIDOTE–AMPHIBOLITE 307
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peak(M2)toretrogression(M3)wasidentifiedfrom
the kyanite–staurolite-bearing rocks in the FPM
(Fig.2).NoFe–Mgexchangegeothermobarometerisapplicable for the observed mineral assemblage.
Hence,metamorphic P–Tconditionsforeachmeta-
morphicstageareestimatedbasedonavailablepet-rogenetic grids and phase equilibria. In this study,KM2
Ca
Na05 050 50
100100 0
KCa
Na05 050 50
100100100 0
0FEA
LMGMMG
M3
FEALMGMMGinclusions in St
Margarite-bearing metamorphic rocks, Central Alps (Frey et al., 1982)
White micas in metamorphic rocks (Guidoti, 1984)Corundum rocks, Westland, New Zealand (Grapes & Palmer, 1996)
Fig. 6.K-Na-Caplotsofcompositionsofwhitemicasfrom
M2andM3stages.CompositionalrangesofwhitemicasbyFrey
etal.(1982),Guidotti,1984)andGrapes&Palmer(1996)are
alsoillustrated.
16.517.017.518.0
3.0 3.5 4.0 4.5 5.0(Fe+Mg+Mn+Zn) 3+1.5xAl18-xSi8O44FEA
LMGSi + Al + Ti - 8
Fe + Mg + Mn + Zn
Fig. 7.Compositionalvariationofanalyzedstaurolitefrom
FEAandLMG.Thedashedlinerepresentsasubstitutionline
(Mg+Fe)3+1.5xAl18-xSi8O46ofmagnesianstaurolitefrom
theSambagawabelt(Yokoyama&Goto,1987).0.00.51.0
6.0 6.5 7.0 7.50.51.0
0.00.51.01.52.0
0.000.15
0.5 1.0 1.5 2.00.00.5
Si
[6]Al[6]Al + Fe3++T i XMg[A]Na +[A]KX[M4]Na TiFEA
LMG
MMG
FEALMGMMG
FEALMG
MMG
FEA
LMGMMG
FEALMGMMG
Fig. 5.ChemicalcompositionsofCa-amphiboleonvarious
plots(seetextfordetails).308 T. TSUJIMORI & J. G. LIOU
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calculationstoobtainphaseequilibriawerecarriedout
usingversion1.1ofthesoftware BAYESBAYESwithaninter-
nally consistent thermodynamic dataset (Chatterjee
etal.,1998).
M1: relict granulite-facies stage
TheM1assemblageispreservedeitherasrelictAl-rich
diopsideoraspseudomorphsafterspinelandplagio-
clase. Al-rich diopside is an index mineral in lowercrustal basic and ultrabasic granulites with the
Cpx+Opx+Pl+Spl assemblage (e.g. Rivalenti
etal., 1981; Ishiwatari, 1985; Wilshire etal., 1991;
McGuire,1994).AlthoughnorelictplagioclaseoccursintheMMG,thepresenceofrelictplagioclasewithin
M2kyaniteoftheFEAsuggestsitsoccurrencepriorto
kyanitecrystallization.OrthopyroxenemayalsohaveoriginallybeenpresentintheM1,perhapshavingbeen
consumedinsubsequentmetamorphicoverprints.The
Al-rich Cpx+Spl (pseudomorph)+Pl (pseudo-morph)assemblagesuggeststhemedium- Pgranulite
facies(0.8–1.3GPaat>800 /C176C),whichisbracketed
bythespinel-gabbro fieldinthesystem CaO-MgO-Al
2O3-SiO2(Gasparik,1984;Schma ¨dicker,2000).In
thespinel–gabbrofield,theAlcontent(Ca-tscherma-
kitecomponent)ofclinopyroxeneishigherthanthatin
the olivine–gabbro and garnet–gabbro fields, and isstronglyT-dependent(e.g. Obata, 1976; Gasparik,
1984). The Al content (up to 8.5 wt.% Al
2O3)o f
relictclinopyroxenesuggestsatemperatureofaround800–900 /C176C.
M2: kyanite–staurolite-bearing HP stage
The kyanite–staurolite-bearing M2 assemblage,
Hbl+Czo+Ky±St+Pg+Rt±Ab±Crn,characterizesthepeakHPmetamorphismofallFPMrocktypes.SincetheFEAandMMGcontaincorun-
duminsteadofquartz,theequilibriumconditionswere
described byamodelsystem projectedfrom corun-dum. In the model CASH (CaO-Al
2O3-SiO2-H2O)
systemwithexcessCrn+H 2O,theP–Tlimitofthe
Ky+Czostabilityisdefinedbythefollowingreac-tions(Fig.8):
Dsp¼CrnþH
2O; ð1Þ
4AnþCrnþH2O¼2Kyþ2Czo ; ð2Þ
and
4Mrg ¼2Kyþ2Czo þ3Crn þ3H2O:ð3Þ
Thepresenceofparagonite,stauroliteandalbitepro-
videsmoreP–Tconstraints;thestabilityfieldsofthese
phasesareboundedbytheNASH(Na 2O-ASH)and
FASH(FeO-ASH)reactionslistedbelow:
AbþCrnþH2O¼Pg ; ð4Þ
Jd50ðOmpÞþKyþH2O¼Pg ; ð5Þ2Fe-Cld þ4KyþCrn¼Fe-St þ3H2O;ð6Þ
and
JdþQtz¼Ab : ð7Þ
Therighthandsidesofthesereactionsarestableinthe
M2.Thestabilityfieldsshifttowardsthelower- Pside
andlower-Tsidewithdecreasingactivitiesofclino-
zoisiteandstaurolite,respectively(Fig.8).The P–T
condition calculated using the software BAYES
(Chatterjeeetal.,1998)withactivitiesofmineralslis-
tedinTable3is600–850 /C176Cand0.95–1.90GPa.
Ontheotherhand,inclusionofrelictplagioclasein
M2kyanitehasprobablyremainedfromthereaction
(2)duringcoolingwithhydration;thisreactioncon-
strainsaP–TlimitfortheM2to550–800 /C176Cand1.1–
1.9GPaaccordingtotheBAYEScalculation(Fig.8).
The compositional gap between M2 paragoniteand
M2 muscovite yields a temperature of c.600 /C176C
(Blencoeetal.,1994).TheMrg-richparagonitewithin
staurolite may suggest a temperature of >600 /C176C
(Franz,1977).ThehighAl(upto18wt.%Al 2O3)and
moderateNa(upto3.4wt.%)contentsofhornblende
further support the HP condition for these rocks
(Ernst&Liu,1998;Niida&Green,1999).
M3: decompression stage
TheM3isdocumentedbyminorretrogradecoronasarmouredaroundM2kyanite,andischaracterizedby
theMrg+Pg±Pl+Chlassemblage.Thisassem-
blageisabreakdownproductofKy+Czoaccordingtothefollowingreactions(Fig.8):
2Kyþ2Czo ¼Mrgþ3An ; ð8Þ
and
2Kyþ2Czo þAb¼Pgþ4An : ð9Þ
Asbothreactions(8)and(9)havegentlypositive P–T
slopes,theM3assemblagemustinvolveacomponent
ofdecompression;its P–Ttrajectorymustcrossthese
reaction curves. Staurolite was not replaced by theCld+Ky assemblage during retrogression; hence
reaction(6)alsoconstrainsthelowertemperaturelimit
to450–500 /C176Catamaximumpressureof<0.5GPa
fortheM3usingthe
BAYESBAYEScalculation(Fig.8).The
Pg–MrggapoftheM3micasuggestsatemperatureof
c.480 /C176C(Blencoeetal.,1994).
DISCUSSION AND CONCLUSION
TheFPMrockspreserveminorrelictmiddle- Pgran-
ulite facies (M1) assemblage of Cpx+Pl+
Spl±possible Opx. They have significantly recrys-
tallized under the HP epidote–amphibolite faciesmetamorphism(M2)closetotheeclogitefaciesand
formed the Hbl+Czo+Ky±St+Pg+Rt±
Ab±Crnassemblage(M2).Subsequentdecompres-sion produced the Mrg+Pg±Ms±Pl+ChlKYANITE–STAUROLITE-BEARING EPIDOTE–AMPHIBOLITE 309
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assemblage(M3).The P–Tpathforthesethreemeta-
morphiceventsisshowninFig.8.
Theestimated P–Tconditionsofthe M2implya
geothermalgradientaround15 /C176Ckm)1,indicatinga
subductionzonemetamorphism. Thedecompression
pathfromM2toM3isroughlyisothermalandsimilartosomecollision-typeHP-UHProcks(e.g.Maruyama
etal.,1996).TheM3mayrepresentagreenschistfacies
overprintduringexhumation;hornblendeK ⁄Arages
of443–403Ma(Tsujimori etal.,2000)maysignifythe
timingofexhumationduringthetrajectoryM2–M3.
ThetimingoftheM1hasnotbeendated,butitmustbeolderthantheK ⁄Arages.
Thepresenceofmedium- Pgranulite-facies(M1)and
thecoolingtrajectorytotheHPmetamorphism(M2)is
anunusualfeature,particularlyinaPacific-typeoro-gen.Lowercrustalgranulitesthatwererecrystallized
andhydratedunderHPorUHPmetamorphismoccur
inmanycontinentalcollisionzones(e.g.GilIbarguchietal.,1991;Tenthorey etal.,1996;Mu¨ntteneret al.,
2000),becausegranulitefaciesrocksconstitutemajor
Precambrian continental crust. However, subducted/C212allochthonous /C213granulitesarelesscommoninPacific-
type orogens. The relict granulite facies assemblage
(Cpx+Opx+Pl+Spl±Grt)intheeclogiteunitofCretaceousSambagawabelt,Japanisarareexam-
ple.There,therelictgranulitesofbasictoultrabasic
cumulateoriginwererecrystallizedundertheCreta-ceousSambagawaHPmetamorphismtoformunusual
assemblages of Ky+Zo±Grt and rare
Ky+St+Crn (Yokoyama, 1980; Yokoyama &Goto,1987).ThemetamorphicevolutionoftheFPM
describedaboveissimilartothatoftheSambagawa
/C212granulites /C213.
Medium-Pgranulitefaciesconditionsareattainedat
Moho depth beneath large oceanic plateaux (e.g.
Saundersetal.,1996).Infact,rareoceanictwo-pyrox-
enegranuliteshavebeenaccretedasaminorcomponentofophiolitetothecircum-Pacificorogen(e.g.Yakuno
(SWJapan):Ishiwatari(1985);Tonshina(Alaska):De-
Bari&Coleman(1989);Bikin(FarEastRussia):Vy-sotskiy,1994);theyhavealsobeendescribedaslower
crustal xenoliths in the present-day oceanic plateau
(Gregoireetal.,1994).Thebulk-rockchemistryimpliesFig. 8.P–Tdiagramsshowingselectedphaserelationshipsandaqualitative P–TpathoftheFPM.Allreactionsarecalculatedby
thesoftwareBAYES(Chatterjee etal.,1998).(a)StabilityfieldoftheKy+Czo+St+AbassemblagewithexcessCrnandH 2O
withend-membercompositionisshowninashadedarea.(b)Aqualitative P–TpathoftheFPM.Theselectedreactioncurvesare
calculatedwiththeactivitieslistedinTable3.Thereactionscurvesof(2’)and(6’)representisoplethsofthereactions(2)and(6)in(a),respectively.ThepetrogeneticgriddefiningtheM1conditionisafterSchma ¨dicker(2000).310 T. TSUJIMORI & J. G. LIOU
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thattheprotolithsoftheKy-St-bearingrocksinthe
FPMarecorrelativewithtroctoliticcumulusrocks,and
coincidentwithcumulatesofthe /C212non-islandarcorigin /C213
ophiolites (Tsujimori &Ishiwatari, 2002).Therelictgranulite facies in both FPMand Sambagawa mayprovideconstraintsonthesubductionofthickoceanic
crust.Saunders etal.(1996)arguedthattheintroduct-
ionofH
2O-richfluidatthelowercrustalpartofan
oceanic plateau allows transformation of gabbroic
granulitetoeclogite.Theypointedoutthatthecollision
ofthickoceaniccrustcausesabackwardmigrationofthesubductionzone,andconsequentlyH
2O-richfluid
releasedfromthesubductingslabpromotestransfor-
mation of suprasubduction zone ophiolitic rocks toeclogite.Theunusual P–TevolutionoftheFPMmay
representsubductionofathick-crustaloceanicplateau
andsubsequentexhumation.
Early Palaeozoic HP metamorphic rocks of the
Kurosegawabelt(Maruyama&Ueda,1974)andthe
Oeyamabelt(Tsujimori etal.,2000)provideapetro-
tectonicconstraintfortheearliestsubductioneventintheJapaneseorogen.TheHProcksdescribedinthis
paperoccurinanearlyPalaeozoicsubductionzonewith
geothermalgradientintheorderof15 /C176Ckm
)1.Sucha
relativelyhighgeothermalgradientinthesubduction
zonehasproducedepidote–amphibolitefaciesmeta-
morphicrocks.Ontheotherhand,typicalblueschists-eclogiteswithalow-geothermalgradientaround10 /C176C
km)1have been exhumed only after Devonian-
Carboniferous time during a continuous subduction
of colder oceanic lithosphere (Ueda etal., 1980;
Tsujimori&Itaya,1999;Tsujimori,2002).Suchadif-
ference and many other examples documented else-
whereledtoasuggestionofsecularcoolingofEarthandsubduction-zonegeothermbyMaruyama etal.(1996).
ACKNOWLEDGEMENTS
Thisresearchwassupportedfinanciallyinpartbya
JSPSResearchFellowshipforYoungScientists,andaGrant-in-Aid for JSPS Fellows to the first author.
T. Itaya and A. Ishiwatari are thanked for helpful
discussion. Preparation of this manuscript was sup-
portedbyNSFEAR-0003355.WethankF.SpearandD.Pattisonfortheirconstructivecomments.
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REVIEW ARTICLE
The basement geology of Japan from A to Z
Simon R. Wallis1| Ken Yamaoka1| Hiroshi Mori2| Akira Ishiwatari3|
Kazuhiro Miyazaki4| Hayato Ueda5
1Department of Earth and Planetary Science,
Graduate School of Science, The University of
Tokyo, 7-3-1 Hongo, Bunkyo-ku, Tokyo,113-0033, Japan
2Geology Course, Faculty of Science,
Department of Science, Shinshu University,3-1-1, Asahi, Matsumoto City, 390-8621,
Japan
3Nuclear Regulation Authority, 1-9-9,
Roppongi, Minato-ku, Tokyo, 106-8450, Japan
4Research Institute of Geology and
Geoinformation, The Geological Survey of
Japan, 1-1-1, Umezono, Tsukuba, Ibaraki,
305-8560, Japan
5Department of Geology, Faculty of Science,
Niigata University 8050, Ikarashi 2-no-cho,
Nishi-ku, Niigata, 950-2181, Japan
Correspondence
Simon R. Wallis, Department of Earth and
Planetary Science, Graduate School of Science,
The University of Tokyo, 7-3-1 Hongo,Bunkyo-ku, Tokyo, 113-0033 Japan.
Email: swallis@eps.s.u-tokyo.ac.jpAbstract
The Precambrian and lower Paleozoic units of the Japanese basement such as the Hida
Oki and South Kitakami terranes have geological affinities with the eastern Asia conti-
nent and particularly strong correlation with units of the South China block. There arealso indications from units such as the Hitachi metamorphics of the Abukuma terrane
and blocks in the Maizuru terrane that some material may have been derived from the
North China block. In addition to magmatism, the Japanese region has seen substantial
growth due to tectonic accretion. The accreted units dominantly consist of mudstone
and sandstone derived from the continental margin with lesser amounts of basalticrocks associated with siliceous deep ocean sediments and local limestone. Two main
phases of accretionary activity and related metamorphism are recorded in the Jurassic
Mino –Tanba –Ashio, Chichibu, and North Kitakami terranes and in the Cretaceous to
Neogene Shimanto and Sanbagawa terranes. Other accreted material includes
ophiolitic sequences, e.g. the Yakuno ophiolite of the Maizuru terrane, the Oeyama
ophiolite of the Sangun terrane, and the Hayachine –Miyamori ophiolite of the South
Kitakami terrane, and limestone-capped ocean plateaus such as the Akiyoshi terrane.The ophiolitic units are likely derived from arc and back-arc basin settings. There has
been no continental collision in Japan, meaning the oceanic subduction record is more
complete than in convergent orogens seen in intracontinental settings making this a
good place to study the geological record of accretion. Hokkaido lacks most of the
Paleozoic history recognized in Honshu, Shikoku, Kyushu, and the Ryukyu Islands tothe south and its geology reflects the Cenozoic development of two convergent
domains with volcanic arcs, their approach, and eventual collision. The Hidaka terrane
reveals a cross section through a volcanic arc and the main accretionary complex of
the convergent system is represented by the Sorachi –Yezo terrane.
KEYWORDS
basement terranes, digital map, geology, Japan, summary
1|INTRODUCTION
The basement rocks of Japan record geological events in the eastern
Asian realm since at least the Cambrian. Throughout much of this timeJapan has been located on or close to a convergent plate boundary.
This long history of convergence has resulted in accretion of bothoceanic sedimentary units, such as chert and mudstone, and siliciclastic
deposits derived from the contemporaneous continental margin; major
examples are the Mino –Tanba –Ashio and the Shimanto terranes. These
sedimentary units are interleaved with more-limited sliced remnants ofoceanic plateau and seamounts. Some of these accreted units can be
traced to the north into Sakhalin island and to the south to Taiwan.Received: 22 June 2019 Revised: 16 December 2019 Accepted: 2 January 2020
DOI: 10.1111/iar.12339
Island Arc. 2020;29:e12339. wileyonlinelibrary.com/journal/iar © 2020 John Wiley & Sons Australia, Ltd 1o f3 1
https://doi.org/10.1111/iar.12339
Connections with the basement of the east Asian continental margin
are shown by the presence of domains of intact continental crust.However, the original geological continuity between Japan and east
Asia has been disrupted by the Miocene opening of the Sea of Japan
and direct correlation is uncertain in many cases. In addition to accre-tion, plate convergence in Japan has also been associated with wide-
spread magmatism represented by multiple intrusive bodies and
volcanic deposits, which cross cut and overlie earlier accreted units.The magma formation involves both the recycling of old crust and addi-
tion of new. The above features make Japan one of the premier sites in
the world to study the effects of plate convergence both in the presentand the past, including subduction, accretion and subduction erosion,and the formation of continental crust.
In this contribution, we present a summary of the basement ter-
ranes of Japan. Several previous excellent summaries have been pres-ented including Isozaki et al. (2010), Wakita (2013) and most recently
inthe Geology of Japan (Chapters 1, 2a –2h, and 3: Ehiro, Tsujimori,
Tsukada, & Nuramkhaan, 2016; Ishiwatari et al., 2016; G. Kimura,Hashimoto, Yamaguchi, Kitamura, & Ujiie, 2016; Kojima et al., 2016;
Miyazaki, Ozaki, Saito, & Toshimitsu, 2016; Taira, Ohara, Wallis,
Ishiwatari, & Iryu, 2016; Tatsumi et al., 2016; Ueda, 2016; S. Wallis &Okudaira, 2016). Our purpose here is to provide a comprehensive
summary of the geological basement terranes of Japan, including
many that have not been closely studied and are not widely knownoutside of Japan (or even within). The alphabetical arrangement of theexplanations is a unique feature of our paper. Our reason for choosing
this dictionary-style presentation is to make it easy for interested
parties to find the relevant explanation for a particular geological ter-rane. Not all geological terranes are introduced with the same level of
detail. This is mainly a reflection of the differences in the state of
research on the different terranes, but is also affected by the researchinterests of the authors. We do not attempt to present a unifying tec-
tonic model. However, in the descriptions of individual terranes we do
highlight tectonic settings that have been proposed.
To place the explanatory text in context and show the geographic
distribution of the terranes and units mentioned, we present a sum-
mary map (Figure 1) and a series of three geological maps covering all
the main landmasses of Japan (Figure 2a –c). The geological map is
based on the map presented by Taira et al. (2016), but is more
detailed and includes numerous modifications that better reflect the
current state of knowledge. Taira et al. (2016) is the introductorychapter to the Geology of Japan ; an important English-language
resource where detailed information can be found both on the geo-
logical units of Japan and how this geology relates to other fields ofrelated studies such as seismology.
One of the motivations for preparing this contribution was to
make available an up to date digital summary map of the geology ofJapan. We hope this will be a useful resource for many students ofJapanese geology. To maximize the utility of this map, we have
decided to make an electronic version available in a format where it
can be freely edited. The map can be downloaded from the websitehttp://science.shinshu-u.ac.jp/~mori/BGJ/index.html. Use of this map
should help reduce the amount of valuable research time spentredrawing the same introductory figure in numerous different labora-
tories around Japan and other countries. Researchers and studentsare free to make use of this resource, although we do ask that they
reference this paper as the original source of information. A second
advantage of making the map available online is that we can incorpo-rate new discoveries or corrections as new outcrops are uncovered,
and the results of new research become available. We invite commu-
nications on these aspects to be sent to base.geol.jpn@gmail.com.Updated versions of our map will be made available when appropriate.
The map includes bathymetry, and the online version is available in
different projections.
1.1 |Belts, terranes, and basement
The distinct geological units of Japan on kilometer scales are generally
referred to as belts, or - tai(帯) in Japanese. Some of these units show
elongate distributions, and their out lines can be said to resemble belts;
however, many do not. To avoid any potential confusion about the
implied shape, here we use the word terrane . To qualify as a distinct ter-
rane in our usage, a geological doma in should show clear differences in
geological history when compare d to its neighboring terranes with
emphasis on age of primary formation, pressure –temperature conditions
and age of metamorphism and lithological associations. Some other stud-ies restrict the use of terrane to faul t-bounded domains (Jones, Howell,
Coney, & Monger, 1983). Our more general usage allows us to include
metamorphic domains such as the Ryoke metamorphism as a distinct
terrane although its boundary with neighboring units is in part gradualand non-tectonic. We also include some terranes, such as the Hida Gaien
terrane, which are complex mixtures of units with different characteris-
tics. In this case, one of the main distinguishing features of the terrane isits complexity and variability. For t he sake of brevity and clarity, we have
tried to keep any subdivisions of t erranes to a minimum. However,
where tectonic subdivisions need to be mentioned we use the term unit.
We also use unitto refer to distinct sedimentary sequences or layers.
We hope with some forbearance the variable uses of the terms terrane
andunitshould not cause too much confusion or irritation.
When preparing our basement geology map, we were also faced
with the problem of deciding what geological criteria we should use
to define basement. First, we exclude nearly all plutonic and recent
volcanic rocks. The only exception is in areas where there is no evi-dence for significant crustal rocks before the development of the cur-
rent distribution of volcanic units, such as the Izu Peninsula. In many
cases, we have also excluded Mesozoic and Cenozoic sedimentarybasin fill. This includes some important depositional sequences such
as the Tetori and Kuruma Formations famous for macro fossil remains;
the extensive Izumi Formation to the north of the Median TectonicLine; and the Monobegawa and Ryoseki Formations overlying theChichibu terrane. We also exclude the deposits formed in the Fossa
Magna –a major rift-zone sequence developed during the Miocene
opening of the Sea of Japan (Figure 1). The relatively young geologyof Hokkaido makes such an approach inappropriate in this area. For
instance, the Yezo Group sediments formed in Cretaceous superficial2o f3 1 WALLIS ET AL.
14401738, 2020, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/iar.12339 by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
forearc basins, but have become so closely involved in the subsequent
orogen that it is more appropriate to include them as part of the
basement.
The removal of large parts of cover sequences may give a mis-
leading impression about some aspects of Japanese geology. For
instance, the geology of the island of Kyushu is currently dominatedby the presence of a set of nested calderas including the voluminous
Mt. Aso (Nakada, Yamamoto, & Maeno, 2016). However, there is no
place for this on our map. Also, a cursory glance at our basement mapof Figure 2 suggests that southwest Japan is geologically considerablymore complex than the equivalent area in the northeast. This is at
least in part a reflection of the poorer outcrop in the northeast; signifi-
cant recent surface uplift and erosion have taken place in southwestJapan which has led to the removal of the volcanic cover revealing
much of the basement geology with its complexities.2|GEOLOGICAL TERRANES
2.1 |Abukuma terrane (Ab)
The Abukuma terrane is situated to the northeast of Tokyo and con-
sists mainly of early Cretaceous intrusive rocks with patches of theoriginal metamorphic basement rock. The age, high T/Pmetamorphic
conditions, and close association between igneous activity and meta-
morphism suggest possible links with the Ryoke terrane of southwestJapan. However, the presence of older protoliths and local evidencefor high-pressure metamorphism (Hiroi, Kishi, Nohara, Sato, & Goto,
1998) show the Abukuma terrane has a more complex history. The
metamorphic rocks of the Abukuma terrane can be separated into theGosaisho, Takanuki, and Hitachi units. There is also a set of undated
metamorphosed limestone, pelite, chert, and mafic-ultramafic rocks
FIGURE 1 Summary map of Japan showing the location of the main islands, cities, volcanoes, and numerous geological features mentioned in
the text. The named volcanoes represent all those listed as rank A and a selected number of those listed as rank B by the Japan MeteorologicalAgency (Japan Meteorological Agency, 2005). The inner and outer zones of Southwest Japan refer to the inner and outer parts of the arc in
Honshu, Shikoku, and Kyushu with the boundary along the major long-lived tectonic boundary, the median tectonic line (MTL). ISTL, Itoigawa –
Shizuoka Tectonic Line; KCTL, Kashiwazaki –Choshi Tectonic Line; MTL, Median Tectonic Line; TTL, Tanakura Tectonic LineWALLIS ET AL. 3o f3 1
14401738, 2020, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/iar.12339 by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
FIGURE 2 A summary map of the basement geology of Japan. The distinction between basement and cover is not clear-cut and an
explanation of the units that were included is given in the main text. A digital version of this map can be downloaded from http://science.shinshu-
u.ac.jp/~mori/BGJ/index.html. The outline map was constructed using GMT software and the bathymetric data are taken from the NOAA dataset
ETOPO1. The outlines of the Daiichi Kashima and Erimo seamounts represent the 5000 m bathymetric contours. (a –c) The areas covered are
shown in Figure 1 [Correction added on 7 February 2020, after first online publication: Figure 2a has been corrected.]4o f3 1 WALLIS ET AL.
14401738, 2020, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/iar.12339 by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
(Ehiro, Kanisawa, & Taketani, 1989) as isolated roof pendants, of
probable oceanic island origin located in the northwest of the terrane.
The main lithologies of the Gosaisho unit are metamorphosed
mafic and siliceous rocks with lesser amounts of pelitic and calcareous
rocks. The lithological association, presence of Early Jurassic radiolar-ians in meta chert and MORB (mid-ocean ridge basalts) geochemistry
of mafic rocks imply an oceanic crustal origin. In contrast, the
Takanuki unit consists mainly of pelitic-psammitic rocks with smallamounts of calcareous, lateritic, siliceous, and mafic rocks. Pelitic-
psammitic rocks are usually migmatitic, indicating that partial melting
took place during high-grade metamorphism. The boundary betweenthe Gosaisho and Takanuki metamorphic units is marked by graniticintrusions and the presence of lenses of ultramafic rocks.
The main metamorphism of the Gosaisho and Takanuki units is part of
the andalusite-sillimanite type progressive metamorphism as defined byMiyashiro (1961). However, careful petrological studies have shown this
was preceded by a phase of short-lived high-temperature (> 700
/C14C) burialto pressures ≥1 GPa in the kyanite stability field. Subsequent exhumation
and metamorphism took place in the Cret aceous (Hiroi et al., 1998; Hiroi &
Kishi, 1989) with both orogenesis-related regional metamorphism and
intrusion-related contact metamorph ism recognized. This type of history
could be related to ridge –trench interactions (Brown, 1998).
Zircon U –Pb dating of the Gosaisho metamorphic rocks yields a
detrital age range of ~ 520 –450 Ma. An igneous crystallization age of ~
122 Ma of dykes that have undergone deformation with the surround-ing metamorphic rocks gives an older limit on the age of orogenesis.
Similar zircon U –Pb dating of the Takanuki metamorphic rocks yields a
metamorphic age of ~ 112 Ma and detrital ages of 280 –200 Ma. The
contrasting detrital zircon ages suggest the Gosaisho unit formed in adistinct paleogeographic setting to the Takanuki unit.
The Hitachi unit is located in the southernmost part of the
Abukuma terrane (Tagiri, Dunkley, Adachi, Hiroi, & Fanning, 2011)and consists of Early and Late Paleozoic sedimentary, volcanic, and
related intrusive rocks typically metamorphosed to the epidote
FIGURE 2 (Continued)WALLIS ET AL. 5o f3 1
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amphibolite facies. Similar lithologies with similar ages are also recog-
nized in the South Kitakami terrane. The Hitachi unit includes Late Cam-
brian sediments –the oldest known sedimentary strata of the Japanese
archipelago. The associated magmatism is evidence for the existence ofa Cambrian arc, which has also been postulated on the basis of detrital
zircon ages (Isozaki et al., 2010). The large age gap between the Cam-
brian rocks and overlying upper Paleozoic strata can be correlated withthe‘great hiatus ’in Paleozoic stratigraphy seen in the North China block
and may indicate a common paleogeographic provenance.
2.2 |Akiyoshi terrane (Ak)
The Akiyoshi terrane is an unmetamorphosed dominantly carbonate
sequence of Carboniferous to Permian age located in the northern
part of western Honshu. The Akiyoshi carbonate units are famous inJapan for the associated Karst geomorphology and limestone caves
and well-preserved fossils of foraminifera, fusulinids, conodonts,
corals, and other reef-forming biota. The stratigraphy of the Akiyoshi
terrane is also important as a long-term (~ 80 my) record of climaticevents and sea-level changes in the Panthalassa Ocean, which can be
related to the growth and retreat of the Late Palaeozoic Gondwana
ice-sheet (Kanmera & Nishi, 1983; H. Sano & Kanmera, 1988).
In more detail, the Akiyoshi terrane consists of a sequence of lower
Carboniferous basalt overlain by Carboniferous to mid Permian lime-
stone and associated with contemporaneous deep-water chert, siliceoustuff, and terrigenous rocks. The basaltic rocks formed by hotspot-typevolcanism in the equatorial zone of the Panthalassa Ocean (S. Sano,
Hayasaka, & Tazaki, 2000; Tatsumi, Kani, Ishizuka, Maruyama, &
Nishimura, 2000). The carbonate rocks show depositional environmentsranging from shallow-marine atoll facies to deeper redeposited clastic
deposits interpreted as atoll-slope units. The radiolarian and sponge-
FIGURE 2 (Continued)6o f3 1 WALLIS ET AL.
14401738, 2020, 1, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/iar.12339 by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
spicule cherts are interpreted as deep marine deposits (H. Sano &
Kanmera, 1988). The Akiyoshi terrane also contains regions of mélangewith a scaly mudstone matrix containing blocks of limestone up to kilo-
meter size. H. Sano and Kanmera (1991a, 1991b, 1991c, 1991d) suggest
the Akiyoshi seamount became dismembered due to forearc bulge nor-mal faulting and the fragmented succession was then incorporated in
the accretionary complex. The Akiyoshi carbonate units are unconform-
ably overlain by Upper Triassic terrestrial sediments giving a youngerlimit on the age of accretion and emplacement at the Earth's surface.
2.3 |Cenozoic volcanics (CV)
Cenozoic volcanic deposits make up a significant part of Japanese
geology. However, here we only include those areas that are part ofthe on-land sequence and yet show little or no evidence for pre-
existing basement lithologies. The two main areas are around the Izu
Peninsula and in northwest Kyushu.
The Izu Peninsula is the northernmost part of the Izu –Bonin –
Mariana (IBM) arc, which collided with the Honshu arc at around
15 Ma causing a strong bending of the preexisting geological structure(Stern, Fouch, & Klemperer, 2003; Takahashi & Saito, 1997; Taylor,
1992). The oldest part of the IBM arc is around 50 Ma and consists of
high Mg andesite, which is thought to occur as an early high Tmantle
melt shortly after the onset of subduction (Tatsumi et al., 2016). Inthis case, the boninite formation can be related to the onset of sub-
duction of the Pacific plate beneath the Philippine Sea plate. This early
phase of volcanism was followed by tholeiitic to calc-alkaline volca-nism with a pause in activity 25 –15 Ma when the Shikoku Basin
opened. Since about 3 Ma the main volcanism has been tholeiitic.
The volcanism in northwest Kyushu is part of the large domain of
alkaline volcanism that affected the northwest coastline region of the
archipelago and was related to the opening of the Sea of Japan
(Mashima, 2009). The main activity is basaltic and concentrated in theperiod 9 Ma to the present although a smaller amount of felsic volca-nism occurred at around 15 Ma (Sakuyama, 2010). Notable Early Mio-
cene moonstone rhyolites erupted in the early stage of the Sea of
Japan opening in coastal areas, and their chemical compositionsclosely resemble Ethiopian rift-zone rhyolites, except for their lower
Nb contents (Ayalew & Ishiwatari, 2011). This phase of magmatism is
thought to be related to mantle upwelling although the scale of thisevent is disputed (Iwamori, 1991; Mashima, 2009; Sakuyama, 2010).
Hattori (1992) reports the presence of metamorphic xenoliths in the
volcanic rocks, and it is possible that this area is underlain by theNagasaki or Sangun metamorphic rocks.
2.4 |Chichibu terrane (Ch)
The Chichibu terrane is one of the main terranes of the outer zone of
southwest Japan (Figure 2) and can be traced over a distance of1500 km from the Kanto Mountains north of Tokyo through Shikoku
and Kyushu to the Ryukyu Islands in the southwest. Despite its goodlateral continuity and wide distribution, this terrane has been less studied
than either the Sanbagawa terrane to the north or the Shimanto terraneto the south. Part of the reason for this is the pervasive low-grade meta-
morphism which has been difficult to a nalyze by traditional petrological
methods but is now receiving more attention (Endo & Wallis, 2017). TheChichibu terrane is commonly divided into northern, middle, and south-
ern units, also referred to as the Chihibu, Kurosegawa, and Sanbosan
belts or terranes, in the same order .H e r ew es u m m a r i z ei n f o r m a t i o nf o r
the northern and southern Chichibu uni ts. The lithologically distinct and
older Kurosegawa terrane is treated as a separate terrane.
2.4.1 |Stratigraphy and age
The Chichibu terrane consists of ocean plate sequence lithologies
including pelagic deposits such as chert and terrigenous sediments
consisting of mudstone and sandstone. Oceanic rock units consisting
of limestone and basalt are also widely developed. Mélange is com-mon and consists of a mudstone matrix enclosing blocks of sandstone,
chert, limestone, and basalt. Microfossils (radiolarians, conodonts,
foraminifera) show sedimentary ages of Late Permian to early Creta-ceous (Matsuoka, Yamakita, Sakakibara, & Hisada, 1998), consistent
with detrital zircon U –Pb ages (Endo & Wallis, 2017; Tominaga,
Hara, & Tokiwa, 2019). These age constraints are important in dis-tinguishing sedimentary units of the Chichibu terrane from theShimanto terrane. The northern Chichibu unit lacks Upper Jurassic
rocks and the southern Chichibu terrane lacks Lower Jurassic rocks
perhaps due to the effects of tectonic erosion. Late Carboniferous toPermian blocks are found in both the southern and northern units.
Younger Triassic limestone is also locally present in the southern unit.
This limestone yields fossils of the bivalve megalodont indicatingdeposition in the Tethyan Realm (Tamura, 1987). Jurassic to Lower
Cretaceous shallow marine limestone referred to as Torinosu lime-
stone is a characteristic deposit of the southern Chichibu terrane andmay have formed on a forearc ridge (Matsuoka, 1992). The presenceof unconformably overlying Lower Cretaceous shallow marine sedi-
ments of the Ryoseki and Monobegawa Groups shows accretion and
exhumation of the Chichibu rocks was complete by this time.
2.4.2 |Structure
The northern Chichibu unit can be divided into several distinct
tectonostratigraphic domains. The structurally lower domains arestratigraphically younger suggesting progressive tectonic stacking
related to subduction accretion. The position of the Chichibu terrane in
the geological framework of southwest Japan is contentious: someworkers suggest it is a klippe (Isozaki, 1996) while others propose it ispart of a strike-slip domain related to large-scale lateral movement
along the Kurosegawa terrane (see discussion in Kojima et al., 2016).
Strong lithological similarities between the northern and southern unitsare evidence in favor of a klippe structure. However, there are also dis-
tinct differences. Two main examples are: the deposition of siliceousWALLIS ET AL. 7o f3 1
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mudstone in the northern unit lasted two to three times longer than
that in the southern unit; and Late Jurassic accretionary domains arecommon in the southern unit but missing in the northern unit. Endo
and Wallis (2017) present direct structural evidence in favor of an over-
all gently dipping klippe structure for the northern unit. However, asimilar detailed study has not been carried out in the southern unit.
2.4.3 |Metamorphism
The northern Chichibu unit has undergone widespread high P/Tmeta-
morphism which can be related to convergent tectonics in a subductionsetting. The progressive northward increase in metamorphic grade from
the Chichibu to the Sanbagawa terranes suggests it is the same meta-
morphism seen in both terranes. However, some old Ar ages (Itaya,Tsujimori, & Liou, 2011) suggest there was also an earlier separate phase
that can be referred to as the Chichibu metamorphism. The peak pres-
sures and temperatures in the northern Chichibu unit of Shikoku rangefrom 0.65 GPa, 290
/C14C close to the Sanbagawa terrane to 0.35 GPa,
220/C14C in the structurally higher southern part. Endo and Wallis (2017)
show that the structural pile is insufficient to account for the differencesin pressure seen and infer there has been significant post metamorphic
tectonic thinning along low-angle normal faults. Similar metamorphic
studies have not been carried out in the southern Chichibu unit.
Metamorphic and structural studies of the Chichibu terrane are
important because they reflect the conditions close to the lower end
of the seismogenic zone in a subduction zone and studies of the asso-
ciated mineral reactions can help identify geological changes that con-trol the change from seismic to non-seismic behavior. Endo and Wallis
(2017) suggest that the breakdown of laumontite may be an important
fluid-releasing reaction that can induce seismicity in subducting slabs.
2.5 |Hida Gaien terrane (HG)
The Hida Gaien terrane occupies a narrow domain located in central
Japan that forms the border region between the Hida –Oki terrane and
a series of neighboring terranes (Figure 2a,b). The definition of this ter-rane differs between different workers (Kojima, Takeuchi, & Tsukada,
2005) and is also referred to as the Hida Marginal belt (or terrane) and
inthe Geology of Japan this terrane is discussed in chapters concerning
both the Pre-Cretaceous accretionary complexes (Kojima et al., 2016) and
Paleozoic basement and associated cover (Ehiro et al., 2016).
The Hida Gaien terrane consists of a series of fault-bounded
mainly Paleozoic sedimentary blocks surrounded by serpentinite
(Chihara & Komatsu, 1982; Komatsu, Ujihara, & Cihara, 1985;
Sohma & Kunugiza, 1993). The structurally lowest part of the strati-graphic section consists of mafic volcanics, followed by Ordovician toCarboniferous felsic tuff and limestone. Mafic rocks become common
in the Upper Carboniferous. These units are followed by Lower to
Middle Permian clastic and pyroclastic rocks. Age constraints on thelower Paleozoic stratigraphy are provided by microfossil finds includ-
ing conodonts, ostracods, and recently chitinozoans (Siveter, Tanaka,Williams, & Männik, 2019; Tsukada, Takeuchi, & Kojima, 2004; Van-
denbroucke et al., 2019) –despite apparently suitable lithologies grap-
tolites have never been reported from this or any other area of Japan.
The Ordovician sediments are the oldest fossiliferous deposits of
Japan. The chronology based on paleontology is supported by zirconU–Pb ages of around 470 Ma (Early to Mid Ordovician) (Nakama,
Hirata, Otoh, & Maruyama, 2010) and 420 Ma (Late Silurian)
(Manchuk, Horie, & Tsukada, 2013) from tuff deposits.
The Hida Gaien terrane has undergone strong tectonic disruption.
Dating of igneous intrusions cross-cutting sheared zones combined
with field observations of deformation features reveal multiple stagesof deformation in the terrane continuing at least until Late Creta-ceous. The presence of 250 –240 Ma granite intruding into both the
Hida Gaien terrane and the neighboring Hida Gneiss unit suggests the
two were juxtaposed by Triassic times. The region is locally overlainby the Jurassic –Cretaceous Kuruma and Tetori Formations.
2.6 |Hidaka terrane (HK) including the Hidaka
metamorphic rocks (HK(m))
The Hidaka terrane is a Paleogene accretionary complex that diagonally
bisects Hokkaido. The terrane has a maximum of 100 km wide but
thins to less than half the distance toward the south. The dominantlithologies are turbiditic siliciclastic rocks including mudstone and sand-stone and lesser amounts of conglomerate. These are associated with
green and red mudstones, chert, and basalt. There are also chaotic mél-
ange facies consisting of a sheared mudstone matrix and blocks of chertand limestone. Some of the basalt outcrops show evidence for syn-
sedimenary intrusion and may have formed close to an active spreading
ridge. Paleo magnetic data indicate that there has been large-angleclockwise rotation of the sedimentary units and a large part of the tur-
bidite sequence may have originated as trench deposits at the junction
of the Kuril and northeast Japan arcs (Nanayama, Kanamatsu, &Fujiwara, 1993). The accretionary units of the Hidaka terrane are locallyintruded by granite bodies that show contact metamorphism. Unde-
formed biotite hornfels of the contact metamorphism grades into bio-
tite schist of the Hidaka metamorphic rocks to the west.
2.6.1 |Hidaka metamorphism
The original accretion-related structure of the Hidaka terrane has
been overprinted by the effects of collision between two convergentplate margins: the northeast Japan and Kuril arcs (Arita et al., 1998;
G. Kimura, 1986). The deeper parts of this collision zone are exposed
in the southwestern part of the Hidaka terrane and include a largedomain of high-temperature and low-pressure metamorphic rocksunderlain by the Horoman peridotite body (Ozawa & Takahashi,
1995). The Hidaka metamorphic rocks form a domain trending NNW –
SSE with a length of about 120 km and a width of 10 –20 km. The
base of this domain is defined by the east-dipping Hidaka Main
Thrust. The foliation generally shows a similar eastward dip.8o f3 1 WALLIS ET AL.
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The grade of the Hidaka metamorphism increases with structural
depth. The structurally upper greenschist and amphibolite parts are dom-inated by pelitic rocks and are commonly associated with granite intru-
sions. The metamorphic rocks of the underlying higher-grade parts are
dominated by amphibolite with less er amounts of pelitic gneiss, and
these are commonly associated with gabbro or tonalite intrusions
(Maeda & Kagami, 1996). The highest-grade garnet- and orthopyroxene-
bearing granulitic rocks are only found as xenoliths. Their presence sug-gests lower crustal metamorphic rocks are present at greater depths than
those currently exposed (Shimur a, Owada, Osanai, & Kagawa, 2004).
The overall trends in metamorphic grade are broadly compatible with asingle phase of regional metamorphism but U –Pb geochronology (Kemp
et al., 2007; Usuki, Kaiden, Misawa, & Shiraishi, 2006) suggests the pres-
ence of two separate phases of metamorphism and igneous activity at
40–30 Ma and ~ 19 Ma. In many cases, a clear distinction between
these two stages is difficult in the fi eld, and it seems similar processes
occurred in the same place separated by 10 –20 my.
The Hidaka terrane represents one of the most accessible sec-
tions through arc crust anywhere in the world with a total of
22–23 km structural section exposed (Komatsu, Miyashita, Maeda,
Osanai, & Toyoshima, 1989; Komatsu, Toyoshima, Osanai, & Arai,1994). The granulite metamorphism is an important example of young
high Tmetamorphism (Kemp et al., 2007), and the Horoman body is
an unusually fresh kilometer-scale peridotite that has been the sourceof valuable information about upper mantle processes.
2.7 |Hida–Oki terrane (HO)
The Hida –Oki terrane lies on the Sea of Japan coastal side including
the Island of Oki and consists dominantly of a granite-gneiss withmigmatite, impure marble, amphibolite, and minor aluminous pelitic
schist. The grade of metamorphism is generally amphibolite to granulite
facies. The Hida gneiss complex has been subdivided into ‘inner ’lower
temperature and ‘outer ’higher temperature regions, based on the min-
eralogy of pelitic gneiss (M. Suzuki, Nakazawa, & Osakabe, 1989).
The Unazuki metamorphic rocks occupy a distinct region of
Barrovian-type metamorphism separated from the main body of Hidagneiss by a mylonite zone. The Unazuki metamorphic rocks include mafic,
quartzofeldspathic, pelitic, and rare rudaceous schists associated with
impure marble. The Unazuki pelitic s chist is well known for the develop-
ment of centimeter-sized well-formed staurolite crystals –a rarity in Japan.
Zircon U –Pb geochronology in metamorphic and igneous rocks indi-
cates the main phase of Hida –Oki metamorphism and magmatism took
place at around 250 Ma close to the Permo-Triassic boundary. Fossils of
bryozoa and foraminifera give a sedimentary age of late Carboniferous
for limestone of the Unazuki schist (Hiroi, Fuji, & Okumura, 1978). How-ever, crystallization ages of around 1.9 Ga (Cho, Takahashi, Yi, & Lee,2012) suggest a much older origin for some of the Oki gneiss. The pelitic
schist of the Hida gneiss yields detrital zircon U –Pb ages of ~ 1.84 Ga
(Y. Sano, Hidaka, Terada, Shimizu, & Suzuki, 2000).
The lithological associations of the Hida –Oki terrane, and particu-
larly the presence of aluminous mudstone, felsic volcanic rocks, andabundant impure siliceous limestone associated with granitic gneiss
suggest a passive-margin origin (Isozaki, 1996, 1997; Wakita, 2013) andvarious possible links with geological units of eastern Asia have been
proposed. Zircon ages show the Hida gneiss and Unazuki granite con-
tain Archean and Paleoproterozoic components but lack evidence forNeoproterozoic zircons. These characteristics suggest a correlation with
the North China block (Horie et al., 2010; Y. Sano, Hidaka, et al., 2000).
The evidence for metamorphism and magmatism around 250 Ma sug-gest a relationship with the Sulu –D a b i eo r o g e nf o r m e db yc o l l i s i o n
between the North China and South China blocks. There may also be
link with the Ogcheon Belt in the Korean Peninsula, where similarlithologies to the Hida –Oki terrane have been reported. Isozaki et al.
(2010) suggest the Unazuki schists may correlate with both the Higo
unit of the Ryoke terrane and the Hitachi unit of the Abukuma terrane
and that all represent fragments of the Sulu –Dabie collisional zone.
2.8 |Idonnappu terrane (including the Poroshiri
ophiolite) (Id)
The Idonnappu terrane lies parallel to and east of the Hidaka terrane in
central Hokkaido and is closely associated with the Sorachi –Yezo ter-
rane to the west (Iwasaki, Watanabe, Itaya, Yamazaki, & Takigami,
1995; Watanabe, Iwaki, Ueda, & Koitabashi, 1994). The Idonnappu ter-rane has an overall west-dipping structure and consists of the Naizawa,Oku-Niikappu, and Horobetsugawa units (Ueda, Kawamura, & Iwata,
2001; Ueda, Kawamura, & Niida, 2000; Ueda & Miyashita, 2005). The
structurally lower units in the east are younger, suggesting progressiveaccretion from west to east (Niida & Kito, 1986).
The Naizawa unit is dominated by basalt, Permian –Triassic lime-
stone, and Triassic to earliest Cretaceous chert, associated with hem-ipelagic mudstone and terrigenous clastic rocks of middle Early
Cretaceous ages (Kato, Iwata, Uozumi, & Nakamura, 1986; Kiyokawa,
1992; Sakakibara, Ofuka, et al., 1997; Sakakibara, Hori, Ikeda, &Umeki, 1997; Ueda et al., 2001). Ueda et al. (2000) identified twokinds of basalts of differing occurrence and chemistry and interpreted
the unit as a tectonic mixture of subducted Triassic seamounts with
the Horokonai ophiolite that makes up part of the Sorachi Group.
The Oku-Niikappu unit is an ophiolitic mélange consisting of
slices and blocks of serpentinized peridotite, gabbro-diorite, altered
volcanic rocks, earliest Cretaceous chert, and mid-Cretaceous clasticsedimentary rocks separated by sheared serpentinites. All the igneous
rocks show island arc chemical characteristics. The island arc volcanic
and plutonic rocks also occur as sedimentary clasts in conglomeratesinterbedded with pelagic chert, which conformably overlie andesite
and boninite volcanics. Ueda and Miyashita (2005) considered that
these rocks were accreted fragments of a subducted intra-oceanicremnant arc like present-day Kyushu-Palau Ridge, where deep crustal(and potentially upper mantle) sections of an island arc had been
exhumed in a pelagic environment owing to back-arc spreading off-
shore. The Horobetsugawa unit mainly consists of Late Cretaceousclastic sedimentary rocks tectonically intercalated with thin sheets of
MORB-type mafic rocks and chert with pelagic red mudstone of mid-WALLIS ET AL. 9o f3 1
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Cretaceous age (Kiyokawa, 1992; Ueda et al., 2000, 2001). This unit
also locally contains sedimentary mélange with blocks probablyderived from the Naizawa unit, and a Palaeocene clastic sequence
that probably represents trench slope basin deposits. The structurally
lowermost parts of this unit underwent greenschist facies metamor-phism. The easternmost part of the Idonnappu terrane is overturned
and overthrust by the Hidaka metamorphic rocks.
2.8.1 |Poroshiri ophiolite
The easternmost part of the Idonnappu terrane is occupied by the
N–S trending Poroshiri ophiolite (Miyashita, Kizaki, Arai, & Toyoshima,
1994), which consists mainly of mafic volcanic rocks and gabbro and
lesser amounts of peridotite and sedimentary rocks. The ophioliteshows widespread greenschist to amphibolite facies metamorphism
with local granulite facies assemblages also recognized. Thrusting and
folding have caused considerable disruption of the original ophiolitesequence (Miyashita, 1983). This deformation has also caused inter-
leaving of the Hidaka metamorphic rocks with the ophiolite (T. Arai &
Miyashita, 1994; T. Arai, Miyashita, & Shimura, 1995).
The geochemistry of the basalt and mineralogy of cumulate and
peridotite all suggest that the ophiolite originated from normal oce-
anic crust formed at an oceanic spreading ridge (Miyashita, Adachi,Tanaka, Nakagawa, & Kimura, 2007; Miyashita et al., 1994). Zircon U –
Pb dating of plagiogranite yields an age of 96.7 Ma ±2.6 Ma (Kizaki,
2000) contemporaneous with pelagic sedimentation in the Late Creta-
ceous unit of the Idonnappu terrane. Ueda (2006) suggests thePoroshiri ophiolite represents part of the subducted oceanic slab on
which the Idonnappu accretionary complex was formed.
2.9 |Joetsu terrane (Jo)
The Joetsu terrane lies immediately to the east of the Fossa Magna
and adjacent to the Mino –Tanba –Ashio terrane. The geology of the
Joetsu terrane is complex, but the rock types can be broadly divided
into three different associations (Takenouchi & Takahashi, 2002).
1. Low grade metamorphic rocks that consist mainly of banded schistose
pelitic rock interleaved with serpentinized harzburgite and dunite asso-ciated with metamorphosed gabbro and basalt. The lithological associ-
ations and 310 –280 Ma K –Ar ages of the metamorphic rocks suggest
possible relationships with the Renge high P/Tmetamorphic rocks of
the Sangun terrane or the Yakuno ophiolite of the Maizuru terrane.
2. Permian units that consist mainly of black shale that may correlate
with the Ultra Tanba terrane based on the presence of Late Perm-ian radiolarian fossils (H. Suzuki & Kuwahara, 2003). There is also aseries of mélange and coherent units with chert layers and blocks
of Permian to Jurassic age that can be correlated with the Mino –
Tanba –Ashio terrane.
3. Shallow marine and brackish water deposits of Permian to Upper
Triassic age that can in part be correlated with the Maizuru terrane(F. Kobayashi, 1955). Some may also correlate with the Tetori and
Kuruma Formations. Associated limestone blocks contain Permianfusulinids.
Further refinement of our understanding of this area is likely to
show that the Joetsu terrane is in fact an amalgamation of several
independent terranes.
2.10 |Kurosegawa terrane (Ks)
The Kurosegawa terrane is a narrow generally E –W trending domain rec-
ognized from Kyushu to the Kanto mountains, northeast of Tokyo and lies
between the northern and southern units of the Chichibu terrane. The
Kurosegawa terrane consists of numerou s blocks of diverse lithologies sur-
rounded by a matrix of serpentinite. The serpentinite is derived mainly
f r o md u n i t ea n dh a r z b u r g i t ea n dd i s t r i b u t e da ss h e e t - l i k eb o d i e su pt os e v -
eral kilometers wide (Hada, Ishii, Landis, Aitchison, & Yoshikura, 2001). Themain blocks within the serpentinite con sist of pre-Silurian basement litholo-
gies including biotite gneiss, amphibolite, and garnet-clinopyroxene granu-
lite with radiometric ages of 540 –400 Ma (Osanai et al., 2000; Yoshikura,
Hada, & Isozaki, 1990) overlain by Siluro –Devonian tuff and tuffaceous
mudstone and sandstone (Yoshikura & Sato, 1976). A mid-Silurian age
(/C24430 Ma) for some of the welded tuff was confirmed by zircon U –Pb
SHRIMP dating (Aitchinson, Hada, Ireland, & Yoshikura, 1996). The sedi-mentary sequence also locally includes limestone with Silurian to Devonian
coral, trilobite, and conodont fossil a ssemblages (Männick et al., 2018; Niko,
Hamada, & Yasui, 1989; Stocker et al., 2019). The Kurosegawa terrane alsocontains numerous blocks that are en closed in serpentinite and have
unclear relationships with the basemen t and covering strata (Ichikawa, Ishii,
Nakagawa, Suyari, & Yamashita, 1956). Some of these blocks are mafic andpelitic schists with a range of metamorphic grades including pumpellyite-
actinolite, blueschist, and epidote -amphibolite facies. The assemblage
jadeite-glaucophane is also locally present. K –Ar ages range from 330 to
150 Ma (Maruyama, Ueda, & Banno, 1978; Wakita, Miyazaki, Toshimitsu,Yokoyama, & Nakagawa, 2007). There are also granitoid blocks and iso-
lated shallow marine deposits of Permian age. A younger limit on the timing
of exhumation and exposure of the main units of the terrane is given bythe presence of unconformably overlying Late Triassic shallow marine
strata. There are strong similarities between the many of the units of the
Kurosegawa and South Kitakami terranes and the two are thought to haverelated geological histories (Ehiro, 2000).
2.11 |Maizuru terrane (including the Yakuno
ophiolite) (Mz)
The Maizuru terrane is a late Paleozoic back arc system with a width
of 10 –30 km that extends ~ 400 km in an arc starting in the Oshima
Peninsula of Fukui Prefecture (the Oshima Peninsula in Hokkaido has
the same spelling but a short initial vowel sound and these are distinctwhen written in Japanese) then tracing to the south and southwest
inland before bending to the west and ending again on the Sea of10 of 31 WALLIS ET AL.
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Japan coast of Shimane Prefecture. The Maizuru terrane can be
divided into three geologically distinct units.
The Northern unit mainly consists of variably deformed granite
bodies associated with metagabbro, amphibolite, and garnet –biotite
gneiss. Dating of granitic gneiss has revealed the presence ofkilometer-scale bodies with formation ages of upto 1830 Ma in this
area, representing the oldest basement rock yet reported in Japan
(K. Kimura, Hayasaka, Shibata, Kawaguchi, & Fujiwara, 2019). Archeangrains are also recorded. SHRIMP zircon U –Pb dates of 425 –405 Ma
for deformed granite and 250 –240 Ma for the undeformed granite
have also been reported (Fujii, Hayasaka, & Terada, 2008). This groupof ages can be correlated with events recognized in the eastern mar-gin of the North China Craton (K. Kimura et al., 2019).
The Central unit is derived from a Permian inter or back arc basin.
The main rock types are a metamorphic sequence with basalt andbasaltic tuff, gabbro, and dolerite protoliths associated with siliceous
mudstone overlain by sandstone and mudstone interbedded with tuff
and limestone. The sediments are collectively known as the MaizuruGroup (Nakazawa, 1958; Shimizu, 1962; S. Suzuki, 1987) and yield
Late Permian to Triassic fossil fusulinids, radiolarians, molluscs, and
brachiopods. The presence of 260 Ma and 520 –380 Ma detrital zir-
con and monazite ages (M. Adachi & Suzuki, 1992) is compatible with
derivation of some clastic material from the Northern unit.
The Southern unit is dominated by the Yakuno ophiolite, which
consists of mantle peridotite, cumulate sequences, gabbro, dolerite,and basaltic volcanics associated with black mudstone (Ishiwatari
et al., 2016). This ophiolite formed in an oceanic island arc –back arc
setting in the Permian (Hayasaka, 1990; Ichiyama & Ishiwatari, 2004;Ishiwatari, Ikeda, & Koide, 1990; Suda, 2004). Geochemical studies
and U –Pb zircon geochronology show this igneous complex initially
formed with MORB or oceanic plateau type magmatism in the period~ 340 –320 Ma before changing to island-arc type magmatism in the
period ~ 290 –280 Ma (Herzig, Kimbrough, & Hayasaka, 1997).
The Maizuru terrane is locally unconformably overlain by Triassic
formations composed mainly of fluvial fan sandstone and conglomer-ate with intercalated mudstone.
2.12 |Mino –Tanba –Ashio terrane (MT)
The Mino –Tanba –Ashio terrane is a Jurassic accretionary complex
that structurally underlies the series of Paleozoic terranes to the
north. Associated sequences are developed throughout most of Japan
from the Ryukyu Islands in the southwest to northeast Japan includingHokkaido. Low-grade metamorphism in the prehnite –pumpellyite
facies is widespread. The Cretaceous Ryoke high T/Pmetamorphic
terrane is mainly developed in the southern part of the Mino –Tanba –
Ashio terrane, and there is no clear thermal discontinuity between thetwo domains of metamorphism.
The Mino –Tanba –Ashio terrane shows a typical oceanic plate
stratigraphy with basalt at its base overlain by limestone followed bychert and clastic rocks representing the approach of a plate toward an
active margin. The basaltic rocks are generally massive but locallyshow pillow or brecciated structures. Limestone blocks yield Early to
Middle Permian fusulina, bivalve, brachiopod, coral, and gastropodfossils and overlie basalt with oceanic island geochemical characteris-
tics. Thinner layers of Upper Triassic deep-marine limestones
(5–10 cm in thickness) are interbedded with chert and characteristi-
cally include shells of planktonic bivalves and radiolarian tests. Detrital
materials in the clastic layers were derived from a Precambrian conti-
nent including high-grade metamorphic rocks (M. Adachi & Kojima,1983; Shibata & Adachi, 1974; K. Suzuki, Adachi, & Tanaka, 1991;
Takeuchi, 2000).
There is an overall younging of the units from north to south or
west to east. The younger units in the southern area are in a structur-ally lower position. These relationships reflect subduction accretion
associated with plate motion toward the continental margin of east
Asia. Thrusting is associated with formation of duplex structureswhich have been overprinted by later large-scale folding (K. Kimura &
Hori, 1993; Matsuoka et al., 1994; Yao, Matsuda, & Isozaki, 1980;
S. Yoshida & Wakita, 1999). Major structural discontinuities are com-monly marked by the presence of mélange.
The Mino –Tanba –Ashio terrane is too old for it to be possible
to relate its formation to movement of a particular plate. However,the slow rate of accumulation of deep-sea units during the change
from the Permian to the Triassic make this terrane well suited to
studying paleoenvironmental chan g e sa c r o s st h i si m p o r t a n tg e o l o g i -
cal boundary (H. Sano, Wada, & Naraoka, 2012). Studies of thisboundary have shown evidence for a large meteorite collision at this
time (Onoue et al., 2012). Careful reconstruction of chert deposits
in the Inuyama area, Gifu Prefecture, has provided a continuousrecord of about 70 my of deep sea sedimentation and these show
astronomical influences on sedimentary cycles (M. Ikeda, Tada, &
Sakuma, 2010).
2.13 |Nagasaki terrane (Sanbagawa terrane?) (Na)
The Nagasaki high P/Tmetamorphic terrane outcrops in the west of
Kyushu and consists of pelitic schist with lesser amounts of quartz
and mafic schists. Lenses and blocks of serpentinite are also foundscattered throughout the higher grade regions. Mica Ar and zircon U –
Pb dating shows that sedimentation and metamorphism occurred in
the Cretaceous (Miyazaki, Suga, et al., 2019; Nishimura, Hirota,Shiokai, Nakahara, & Itaya, 2004; Tsutsumi, Horie, Miyashita, &
Shiraishi, 2011). The foliation is gently dipping in all areas and the
stretching lineation is approximately north −south.
There are similarities in lithology, metamorphic conditions, and
ages of both protolith formation and metamorphism with the San-
bagawa terrane of southwest Japan (Miyazaki et al., 2016). However,a correlation between the two terranes is not universally accepteddue to the following differences. (i) If the Nagasaki and Sanbagawa
terranes do correlate, this implies a major clockwise rotation in the
trend of the orogen from roughly east −west in Shikoku to
north−south in the Kyushu region. Such a change in orientation would
be opposite to the anticlockwise changes in direction shown by theWALLIS ET AL. 11 of 31
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Chichibu and Shimanto terranes to the south (Figure 2). (ii) There are
other clear differences such as the tectonic transport direction: in theNagasaki terrane this is roughly north −south whereas in the San-
bagawa terrane it is dominantly east −west to WNW –SSE.
(iii) Another difference is that in areas to the east, the boundary withthe Ryoke terrane, marked by the Median Tectonic Line, is a clear dis-
tinguishing feature of the Sanbagawa terrane, but is lacking in the
region around the Nagasaki terrane. (iv) It may also be significant thatrecent identification of metamorphism under diamond-stable condi-
tions in the Nagasaki terrane (Nishiyama, Nishi, Harada, Ohfuji, &
Fukuba, 2018) has not been reproduced in the Sanbagawa terrane.The abovementioned (i), (ii) and (iii) differences, however, may beexplained by a clockwise rotation of about 50
/C14and right-lateral strike-
slip of southwest Japan against the Nagasaki area during the opening
of the Sea of Japan (Faure, Fabbri, & Monie, 1988). In addition, theYobikonoseto Fault in the Nishisonogi Peninsula and Wakimisaki-
Fukabori Fault in Nomo Peninsula are correlated with the Paleo-
Median Tectonic Line (Nishimura et al., 2004).
The main metamorphism of the Nagasaki terrane is in the
epidote –blueschist to albite –epidote amphibolite subfacies, equiva-
lent to the garnet to albite –biotite zones of Higashino (1990) defined
in Shikoku and there is an overall increase in metamorphic grade
structurally downwards in the Nomo Peninsula (Miyazaki &
Nishiyama, 1989). Higher grade blocks ，such as garnet-bearing
pyroxenite and quartz-bearing jadeitite, are locally present enclosedwithin the serpentinite in the Nishisonogi Peninsula (Nishiyama, 1989;
Shigeno, Mori, & Nishiyama, 2005). Kyanite-bearing amphibolite and
high-pressure granulite facies rocks tectonically overlie epidote-blueschist facies schists in the Amakusa Shimojima area, and the amal-
gamation of these rocks probably took place at lower crustal depths
(Miyazaki et al., 2013).
The Nagasaki metamorphic rocks are in fault contact with overly-
ing distinct domains of Triassic –Jurassic high P/Tmetamorphic rocks
that can be assigned to the Suo unit of the Sangun terrane(Nishimura, 1998; Nishimura et al., 2004) and in places with UpperCretaceous sediments.
2.14 |Nedamo terrane (Nd)
The Nedamo terrane is a Carboniferous accretionary complex, the
oldest such domain in Japan and occupies a relatively small area
between the South Kitakami and North Kitakami terranes (Figure 2).
The main lithologies are mafic rock, chert, interbedded mudstone andfelsic tuff, sandstone, and conglomerate. Mélange or broken forma-
tions are common. The mafic rocks are mainly basaltic volcaniclastics
and massive lava locally with pillow structures. These rocks show bothoceanic island and MORB geochemistries (Hamano et al., 2002;Uchino & Kawamura, 2009). Prehnite-pumpellyite to epidote
blueschist facies parageneses are present in these rock types (Moriya,
1972; Onuki, Shiba, Kagawa, & Hori, 1988; Uchino & Kawamura,2010a, 2010b). Chert shows centimeter-scale iron-manganese layers
and yields Middle to Late Devonian radiolarians, conodonts, andsponge spicules (Hamano et al., 2002; T. Kawamura et al., 2013). The
chert is overlain by units including large amounts of felsic tuff. Thepresence of this tuff distinguishes the Nedamo terrane from other
accretionary terranes of Japan. Mudstone and the muddy matrix to
conglomerate locally yields Early Carboniferous radiolarians rep-resenting the age of trench-fill sediments shortly before accretion
(Uchino, Kawamura & Kurihara, 2005; Uchino & Kurihara, 2019). Ser-
pentinized mantle-derived rocks and glaucophane-bearing schist(Uchino & Kawamura, 2006) occur as blocks decorating tectonic
boundaries within the terrane and along its margins.
The Nedamo terrane has undergone ductile deformation dis-
playing a northwest –southeast striking and southwest dipping folia-
tion commonly associated with tight folds. The best estimates of the
age of metamorphism of the metamorphic blocks come from Ar dating
that yields ages of around 380 Ma (Kawamaura et al., 2007). TheNedamo terrane may be geologically linked to the Kurosegawa and
Sangun terranes.
2.15 |Nemuro terrane (Nm)
The Nemuro terrane makes up a large part of the eastern corner of
Hokkaido and consists of Late Cretaceous to Early Paleogene clastic
sedimentary rocks associated with volcanic and intrusive rocks(Kiminami, Niida, et al., 1992) and represents a forearc sequence. Thelowest upper Cretaceous units contain lava and tuff breccia as well as
coarse clastic rocks. The volcanic rocks consist of basalt, andesite, and
dacite with both tholeiitic and calc-alkaline chemical affinities(Kiminami, Niida, et al., 1992). Overlying sequences consist of muddy,
sandy, and conglomeratic formations, representing turbidite, bottom
current, and slump deposits (Kiminami, 1975, 1983; Naruse, 2003;Naruse & Ohtsubo, 2011). Upper Cretaceous sections locally yield
megafossils such as ammonoids and inoceramids (Naruse, Maeda, &
Shigeta, 2000). The Nemuro terrane extends to the present-day outerarc of the Lesser Kuril Islands.
2.16 |North Kitakami –Oshima terrane (NK)
The North Kitakami –Oshima terrane is an accretionary complex that
occupies a broad region in northeast Japan and west Hokkaido. Themain rock types are basalt, chert with associated hemipelagic siliceous
mudstone, and terrigenous mud and sandstone. Lenses of Carbonifer-
ous to Triassic limestone are also present. Paleontological evidenceshows the chert and basalt sequences range in age from Late Carbon-
iferous to Jurassic (Ueda et al., 2018). The overlying mud and sand-
stone units are dominantly Middle to Upper Jurassic and locally asyoung as middle Early Cretceous (130 Ma: Ueda et al., 2018). LatePermian (Nakae & Kurihara, 2011) and Late Triassic (Uchino, 2017)
clastic rocks are also known to occur in the western marginal parts.
The North Kitakami terrane can be divided into southwest and north-east units based on sandstone compositions (more K-feldspar rich in
the northeast and more plagioclase-rich in the southwest), the12 of 31 WALLIS ET AL.
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presence in the southwest of characteristic Permian and Carbonifer-
ous marine chert and limestone, and Jurassic coral-bearing limestone.Tight folds with an overall shallowly dipping enveloping surface are
common in the north, whereas more open folds are seen to the south
(the northern North Kitakami Massif) (Sugimoto, 1974).
The North Kitakami –Oshima terrane has strong similarities with the
Chichibu terrane and the two likely represent different parts of the
same original geological domain. A particular characteristic shared byboth terranes is the presence of an Upper Carboniferous sequence
(Ehiro, Yamakita, Takahashi, & Suzuki, 2008) from basalt through alter-
nating dolomite-chert beds to red ribbon chert. The main difference isthe width of these two terranes: the North Kitakami –Oshima terrane
reaches 150 km whereas the Chichibu terrane reaches a maximum of
10 km. This may reflect post-accretion tectonic thinning in southwest
Japan. The southwestern margin of the North Kitakami –Oshima terrane
may correlate with the Ultra-Tanba terrane (Nakae & Kurihara, 2011).
The juxtaposition of the North and South Kitakami terranes was
completed after the deposition of the Jurassic Torinosu-type lime-stone and before a major phase of deformation and low-grade meta-
morphism that affected both terranes. A younger age limit for this
orogenic phase is given by the presence of undeformed mid-Cretaceous sediments (Kanisawa & Ehiro, 1989).
2.17 |Rebun –Kabato terrane (RK)
The Rebun –Kabato terrane is a north −south trending volcanic domain
that lies in the eastern part of Hokkaido and borders the NorthKitakami –Oshima terrane to the west. This terrane represents a volca-
nic arc linked to the subduction that also formed the Sorachi –Yezo and
Idonnappu terranes. The main lithologies are mudstone, volcaniclasticsandstone, sandstone, mudstone, and tuff, associated with formations
dominated by volcanic conglomerate, tuff breccia, pillow lava, and
dykes (I. Ikeda & Komatsu, 1986; Kondo, 1991; Nagao, Akiba, & Omori,1963; M. Nagata, Kito, & Niida, 1986). Volcanic rocks are high-K basaltto andesite, including both tholeiitic and calc-alkaline series, with low-
TiO
2island arc characteristics (I. Ikeda & Komatsu, 1986; Kondo, 1993;
M. Nagata et al., 1986). An early to mid-Cretaceous age for most of theterrane is shown by radiolarian and ammonoid fossils (Kondo, 1991;
Nagao et al., 1963; M. Nagata et al., 1986) and Ar –Ar dating of volcanic
rocks (Kondo, 1991). Clasts of Permian fusulina limestone(W. Hashimoto, Igo, Asakura, Tateno, & Nagase, 1960) and of Jurassic
radiolarian siliceous mudstone (Kondo, 1991) are found in the conglom-
erate deposits and were probably derived from the Jurassic accretion-ary complex of the neighboring North Kitakami –Oshima terrane.
2.18 |Ryoke terrane including the Higo unit (Ry)
T h eR y o k et e r r a n ei sar e g i o no fh i g h T/Pmetamorphism and associated
magmatism located to the north of the Median Tectonic Line andstretching from Kyushu in the west to the Kanto mountains in the east
with a distance of around 1000 km. This terrane has a width of30–50 km and is similar in size to that of the Sierra Nevada and Penin-
sular Ranges batholiths in western North America. For most of its length,the Ryoke terrane can be considered a high Tmetamorphic domain
affecting the southern part of the Mino –Tanba accretionary terrane and
accordingly the dominant protoliths are mudstone, sandstone, and chert.The southern margin of the Ryoke terrane is marked by the unconform-
ably overlying Cretaceous forearc basin deposits of the Onogawa –Izumi
Group or cut by the Median Tectonic Line (S. Wallis & Okudaira, 2016).The western and eastern extensions of the Ryoke terrane are less clear.
Correlations have been proposed with the Higo metamorphic unit in
central Kyushu, high T/Pmetamorphic rocks in northern Kyushu, and
the Abukuma terrane to the east of Tokyo. The protoliths in theseregions of metamorphism are distinct from the Mino –Tanba terrane.
Metamorphism of the Ryoke terrane can be divided into six zones
based on the mineral parageneses of metapelite: the chlorite –biotite,
biotite, muscovite –cordierite, K-feldspar –cordierite, sillimanite –K-
feldspar, and garnet –cordierite zones in order of increasing grade
(S. Wallis & Okudaira, 2016). The chlorite –biotite and muscovite –
cordierite zones locally overlap with contact aureoles, which can make
the separation between regional and contact metamorphism difficult
(Skrzypek et al., 2016). Peak temperatures for the different mineralzones show a progressive increase from around 450
/C14C to over
800/C14C. Pressure estimates are variable but low and show a progres-
sion from < 0.1 to ~ 0.7 GPa with increasing peak metamorphic tem-perature. The inferred metamorphic field gradient is 40 –50
/C14C/km.
Metamorphic P–Tpaths are essentially isobaric (Brown, 1998), but
some stages of nearly isothermal decompression have been suggested
(Kamitomo, Imaoka, & Owada, 2008; Kawakami, 2002).
The high Tmetamorphism is closely associated with granitic and
tonalitic intrusions, and these can be divided into older and younger
types (Koide, 1958). The older units are foliated and were intruded assheets or lenses parallel to the gneissose structure of high-grade
domains. The lack of contact metamorphism around the older intru-
sions reflects the importance of these intrusions as a heat source forthe regional metamorphism (Okudaira 1996). The younger intrusionsare nearly unfoliated and their intrusion postdates the main deforma-
tion of the Ryoke terrane. The conditions of early regional and later
contact metamorphism can be very similar, but microstructural obser-vations of aluminosilicate minerals can help distinguish them
(Y. Adachi & Wallis, 2008; Miyake, Murata, & Morishita, 1992).
The main ductile deformation is associated with a well-developed
foliation and generally east −west oriented stretching subparallel to the
terrane. The sense of shear is not always clear but where documented
is shows west ward-directed shear of the overlying units relative to theunderlying (Y. Adachi & Wallis, 2008). Structures related to N –Se x t e n -
sion and magma intrusion are also reported (Okudaira, Takeshita, &
Hara, 1995) and this period may have been a time of closely relatedeast−west and north −south extension. Close to the Median Tectonic
Line, the Ryoke terrane shows high-strain non-coaxial sinistral defor-
mation forming a mylonitic zone up to 2 km wide (Michibayashi &
Masuda, 1993; Okudaira, Beppu, Yano, Tsuyama, & Ishii, 2009). Thekinematics of deformation is similar to the main high Tdeformation of
the Ryoke terrane, but the deformation close to the Median TectonicWALLIS ET AL. 13 of 31
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Line is generally treated as a separate phase. The recrystallized grain
size decreases toward the fault contact, suggesting locally high shearstresses (Okudaira & Shigematsu, 2012; Takagi, 1986).
Upright open folds that deform the main ductile fabrics are com-
mon particularly in the high-grade zones closer to the Median TectonicLine (Okudaira et al., 2001). Locally it can be shown that intrusion of
the younger igneous units began synchronously with and continued
after this stage of folding (Nishiwaki & Okudaira, 2007). Some of theprominent folds in strongly layered meta chert are premetamorphic
and possibly accretion related (Okudaira & Beppu, 2008).
Monazite U –Th–Pb (CHIME) ages suggest the peak of metamor-
phism occurred around 95 –90 Ma (Kawakami & Suzuki, 2011;
K. Suzuki & Adachi, 1998) and the metamorphic and igneous activity
continued until around 70 Ma in the east and 85 Ma in the west.
There is a general eastward younging of K –Ar and Rb –Sr mineral iso-
chron ages and this has been interpreted as the result of oblique sub-
duction of a spreading ridge (Kinoshita, 1995; Kinoshita & Ito, 1986;
Nakajima, 1994, 1996; Nakajima, Shirahase, & Shibata, 1990) or dif-ferential exhumation (K. Suzuki & Adachi, 1998). However, recent zir-
con U –Pb dating reveals crystallization ages younger than some of
those previously reported, suggesting that substantial magmatic activ-ity in the central part of the Ryoke terrane continued longer than gen-
erally thought (Skrzypek et al., 2016; Takatsuka et al., 2018).
Irrespective of such revisions, there is a general overlap in the ageranges of the Ryoke metamorphism and the high P/TSanbagawa ter-
rane to the south and the two have commonly been linked as a pair of
metamorphic belts as originally proposed by Miyashiro (1961).
2.18.1 |Higo metamorphic unit and high T/P
metamorphic rocks in northern Kyushu
The Higo unit is a kilometer-scale met amorphic domain located in west cen-
tral Kyushu with metamorphic rocks showing amphibolite to granulite faciesand migmatite formation at the highest grade. The main protoliths for themetamorphic rocks are sandstone, mud stone, mafic volcanic rocks, lime-
stone, siliceous rock, and ultramafic rocks, which are probably derived from
the Suo unit of the Sangun terrane. Parageneses of the meta-mudstone canbe used to divide the unit into biotite, K-feldspar –sillimanite, garnet –cordi-
erite, and orthopyroxene zones in order of increasing metamorphic grade
(Miyazaki, 2004). Sapphirine-bearing granulite locally occurs as blocksenclosed within metamorphosed ul tramafic rocks (Osanai, Hamamoto,
Maishima, & Kagami, 1998). The Higo unit preserves a sequence about
10 km thick with increasing temperature and pressure down structuralsection (Miyazaki, 2004). Zircon ages (Maki, Fukuyama, Miyazaki, Yui, &
Grove, 2011) suggest the peak of metamorphism occurred 120 –110 Ma
–5–10 my older than the age of other parts of the Ryoke terrane. The
youngest age component of detrital z ircons from pelitic gneisses is ca
200 Ma, and the protolith of the Higo metamorphic unit is probably the
Suo metamorphic unit (Suga et al., 2017) of the Sangun terrane.
Cretaceous high T/Pmetamorphic rocks derived from the Sangun and
Akiyoshi terranes are also present in northern Kyushu. The metamorphic
conditions, metamorphic ages, and overall geographical trend of these unitssuggest a correlation with the Ryoke ter rane is likely (T. Ikeda, Miyazaki, &
Matsuura, 2017; Miyazaki et al., 2017). This domain of high T/Pmetamor-
phism is associated with formation of voluminous igneous rocks synchro-
nous with formation and ascent of migmatite (Miyazaki, Ikeda, et al., 2019).
2.19 |Sangun terrane (Sa) including the Oeyama
ophiolite (Oe), Renge (r), and Suo units
The Sangun terrane is a high P/Tdomain distributed in a broad region
of Honshu and the northern part of Kyushu that includes the prominentOeyama ophiolite unit. It can also be traced further west to the RyukyuIslands. Ar dating of phengite shows that the Sangun terrane actually
consists of two distinct domains: an older Late Paleozoic domain that is
commonly referred to as the Renge terrane or belt and a younger EarlyMesozoic domain that is referred to as the Suo terrane or belt
(Nishimura, 1998; Shibata & Nishimura, 1989; Tsujimori & Itaya, 1999).
However, the two domains overlap in their distributions, have similarprotoliths, and have undergone similar high P/Tmetamorphism. In addi-
tion, the Renge terrane has only limited known exposure. For these rea-
sons, it is impractical to differentiate the two terranes on our map, andhere we show them both as part of a wider Sangun terrane.
2.19.1 |Oeyama ophiolite (Oe)
The Early Palaeozoic Oeyama ophiolite is one of the main units of the
Sangun terrane and occupies the structurally highest position of theaccretionary units of the inner zone of southwest Japan (Ehiro et al.,
2016; Ishiwatari et al., 2016). The ophiolite mainly consists of ser-
pentinized mantle peridotite. Jadeitic and other metamorphic rocksalso occur as blocks in sheared zones. The ophiolite sequence has
undergone contact metamorphism caused by Cretaceous granite
intrusions (S. Arai, 1980; Kurokawa, 1985; Uda, 1984). The originalperidotite was dominantly lherzolite and harzburgite with local duniteand metamorphosed cumulate rocks (Kurokawa, 1985; Tsujimori &
Ishiwatari, 2002). No volcanic components of the Oeyama ophiolite
have been reported. The Oeyama ophiolite is underlain by ser-pentinite mélange with blocks of ~ 300 Ma high P/Tmetamorphic
rocks that can be correlated with the Renge unit (Tsujimori, 1998).
A Cambrian to Ordovician age for the Oeyama ophiolite is
suggested by K –Ar dating of mafic intrusions (~ 450 Ma) (Nishimura,
1998) and zircon U –Pb dating of jadeitic metamorphic blocks (~ 530 Ma)
in the associated mélange (K. Kimura & Hayasaka, 2015). The Oeyamaophiolite is one example of the widesp read ophiolite units formed at this
time, including the Hayachine –Miyamori ophiolite of the South Kitakami
terrane (Ishiwatari, 1991; Ishiwatari, Sokolov, & Vysotskiy, 2003).
2.19.2 |Renge unit
The Renge unit is a high P/Tmetamorphic domain that dominantly
consists of siliceous-mudstone, mudstone, and sandstone associated14 of 31 WALLIS ET AL.
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with lesser amounts of basalt and peridotite. The peridotite may be
associated with the Oeyama ophiolite. The Renge metamorphicrocks occur as tectonic sheets or blocks within serpentinite méla-
nge associated with the Oeyama ophiolite and only have a limited
outcrop distribution. The Renge metamorphism is dominantly in theepidote-blueschist to epidote-amphibolite facies with local exam-
ples of lawsonite-blueschist and glaucophane-eclogite facies
(Tsujimori, 2010).
Zircon U –Pb dating suggests the peak metamorphism occurred
around 360 Ma and Ar-dating of phengite yields an age-range of
360–280 Ma (Tsujimori, 2010). Detrital zircon suggests a Late Devo-
nian protolith age (Tsutsumi, Yokoyama, Terada, & Hidaka, 2011). Theserpentinite mélange associated with the Renge metamorphic rocks is
unconformably overlain by the Kuruma and Tetori Groups showing
exhumation was complete by the early Jurassic. Detrital sedimentarymaterial probably derived from the Renge metamorphic rocks is found
in the Maizuru, Joetsu, and Shimanto terranes suggesting a much
larger original distribution.
2.19.3 |Suo unit
The Suo unit occupies a large part of the inner zone of southwest
Japan and can be traced westwards to Kyushu and the RyukyuIslands. The Suo unit is a metamorphosed Permian to Jurassic accre-tionary complex, consisting mainly of siliceous mudstone, mudstone,
sandstone, and basalt locally associated with limestone, gabbro, and
serpentinite. Rocks equivalent to the Suo unit may be present in theKurosegawa terrane (Isozaki & Itaya, 1991). The age of deposition is
shown by radiolarian and conodont fossils (Hayasaka, 1987; Take-
shita, Watanabe, & Ishiga, 1987).
The main metamorphism of the Suo unit took place in the
pumpellyite-actinolite, epidote-blueschist, greenschist facies, and
epidote-amphibolite facies (M. Hashimoto, 1968; Hayasaka, 1987;Nishimura, 1971; Nishimura et al., 2004). There is also local contactmetamorphism with granite intrusions (Miyakawa, 1961; Shibata &
Nishimura, 1989).
The Suo unit is overlain by the Akiyoshi terrane and is in fault
contact with the Yakuno ophiolite of the Maizuru terrane. Some
workers suggest that the Suo unit in the eastern Chugoku area is the
metamorphosed equivalent of the underlying Mino –Tanba –Ashio
terrane (Hayasaka, 1987). In western Kyushu, the Suo unit is directly
underlain by the Cretaceous schists of the Nagasaki terrane
(Nishimura et al., 2004). Both the Nagasaki and Suo schists showsimilar polyphase ductile deformation (Oho, 1990) and the two are
difficult to distinguish in the field. The main distinguishing feature is
the age. Phengite K –Ar dating yields ~ 220 –190 Ma (Nishimura,
1998; Shibata & Nishimura, 1989) for the Suo schists whereas theNagasaki schists are Cretaceous. U –Pb zircon ages (Miyamoto &
Yanagi, 1996; Tsutsumi, Yokoyama, Terada, & Sano, 2003) support a
Suo subduction metamorphism of around ~ 220 Ma. This is in con-trast with the 120 –70 Ma age of subduction of the Sanbagawa
terrane.2.20 |Sanbagawa terrane (Sb) including the
Mikabu unit (m)
The Sanbagawa terrane is a domain of Cretaceous high P/Tmetamor-
phism located to the south of the Median Tectonic Line. It is a well-studied ancient record of warm subduction and can be closely compared
to modern day warm subduction zones such as southwest Japan and
Cascadia, west coast North America. The timing of the Sanbagawa meta-morphism and kinematic evidence for highly oblique sinistral conver-
gence help identify the subducting oceanic slab as part of the Izanagi
plate (S. R. Wallis et al., 2009). It is contemporaneous with the formationof the Shimanto accretionary complex and may be a more deeply buriedequivalent. To the north, the high T/PRyoke terrane also formed at the
same time and is commonly identifie d as part of the contemporaneous
volcanic arc. Together these two me tamorphic terranes are commonly
referred to as paired metamorphic belts. The main problem with consid-
ering these two domains as contempo raneous genetically linked paired
belts is that they are now adjacent. What has happened to the interven-ing units? In modern settings, the domains of high and low P/Tmetamor-
phism should be separated by 100 –200 km horizontal distance.
2.20.1 |Protolith
The main lithological components of the Sanbagawa terrane are
psammitic, pelitic, quartz, and mafic schists representing a paleo oce-
anic plate stratigraphy of sandstone, mudstone, chert, and basalt. In
central Shikoku, there are also significant units of limestone and meta-gabbro (Aoya et al., 2013; Endo, Wallis, Tsuboi, Aoya, & Uehara, 2012).
Blocks of serpentinized mantle rock are an additional minor but wide-
spread component. These blocks are only present in the high-gradezone implying that the slab-derived units had to be subducted to depth
before they came into contact with the mantle-derived serpentinite
(Aoya, Endo, Mizukami, & Wallis, 2013). This is strong geological evi-dence that the serpentinite is derived from the mantle wedge, and thecontact with the slab-derived units represents a fossil plate boundary.
The Higashiakaishi unit of central Shikoku deserves special mention as
one of the only two known garnet peridotite bodies recognized in anoceanic subduction setting –the others are in continental collision
zones. The Sanbagawa terrane is also well-known in resource geology
as the host of the eponymous examples of Besshi-type Cu deposits.
2.20.2 |Metamorphism
The main Sanbagawa metamorphism can be divided into three mineral
zones based on characteristic Fe –Mg silicates developed in metapelite:
the chlorite, garnet, and biotite zones. The biotite zone can be subdividedinto a lower temperature albite-bearing zone and a higher temperature
oligoclase-bearing zone (Enami, 1983). The degree of metamorphism
shows a clear general increase in gr ade toward the Median Tectonic Line
although folding and faulting locally complicate this trend (Mori & Wallis,
2010). The metamorphic zonation defined by the mineral zones of theWALLIS ET AL. 15 of 31
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main metamorphism cross-cuts an earl ier eclogite facies metamorphism.
There is also local evidence for a pre- eclogite facies metamorphism at rel-
atively high Tand moderate P(Endo et al., 2009). The polyphase meta-
morphic history can be explained as different stages in the thermal
evolution of the same subduction zon e: hot conditions shortly after the
initiation of subduction, deep subduction, and heating during exhumation.
Each of these phases is associated with distinct temperature peaks. The
main phase of subduction-related P–Tpaths shows a distinctive strongly
curved trajectory with increasing P/Tat greater depths.
2.20.3 |Geochronology
The oldest protolith ages from the Sanbagawa terrane are given by
Re–Os dating of Cu-sulphide deposits, which yields an age of 150 Ma
and is interpreted as the result of hydrothermal activity close to a
spreading axis (Nozaki, Kato, & Suzuki, 2013). The next distinct stage
is recorded by zircon U –Pb and garnet Lu –Hf ages, which indicate
mineral growth at 120 –115 Ma (Endo et al., 2009; K. Okamoto et al.,
2004). This can be related to the early stage metamorphism shortly
after the inception of subduction. The best estimate of the peak ofeclogite metamorphism and the main phase of subduction-related
prograde metamorphism is given by Lu –Hf dates of garnet in Shikoku
at around 90 Ma (S. R. Wallis et al., 2009) implying a 30 myr gapbetween initial subduction and the main phase of metamorphismrecorded in the rocks. Detrital zircon ages slightly older than this in
the same area show that psammitic and pelitic units were subducted
shortly after deposition in a trench setting (Aoki et al., 2019; Knittelet al., 2014). Mica Ar –Ar cooling ages combined with P–Tpaths from
central Shikoku suggest exhumation was initially at rates of centime-
ters per year (S. R. Wallis et al., 2009). Zircon fission track and ero-sional ages suggest subsequent exhumation was slower but largely
complete by 55 Ma (S. Wallis, Moriyama, & Tagami, 2004). The order
of magnitude change in exhumation rate suggests a change frombuoyancy-related to erosion-related exhumation.
2.20.4 |Deformation
The Sanbagawa terrane generally has a strong penetrative north-
dipping foliation associated with WNW –ESE to west −east stretching
lineation which formed during exhumation. A younger phase of
upright folding with roughly east −west fold axes is also widely devel-
oped. An earlier polyphase history of ductile deformation can be recog-nized in the high-grade parts of central Shikoku, but this is poorly
preserved in other areas. Subduction- and exhumation-related deforma-
tion stages show plane-strain to flattening finite strain types. Uniaxialextension seen in some rocks is the result of overprinting strain relatedto two separate phases and not a contemporaneous deformation type
(Aoya, Noda, et al., 2013; Moriyama & Wallis, 2002). There is good lat-
eral continuity of many lithological layers, showing that the Sanbagawaterrane can be distinguished from mélange-dominant subduction-type
metamorphic terranes such as Kamuikotan unit of the Sorachi –Yezoterrane. The difference is likely related to the relatively high Tof San-
bagawa terrane. Several distinct nappes can be identified with distinctmetamorphic and structural histories; three commonly identified are the
eclogite, Shirataki or Besshi, and Oboke units (Aoya, Noda, et al., 2013;
S. Wallis & Okudaira, 2016).
2.20.5 |Tectonics
The age of orogenesis in the Sanbagawa terrane and the orogen-
oblique movements are best explained if orogenesis is related to therapid oblique subduction of the former Izanagi Plate with a sinistralsense of shear (S. R. Wallis et al., 2009). The Sanbagawa terrane pre-
serves evidence for events from subduction initiation at around
120 Ma to exhumation at around 60 Ma. The onset of exhumationcoincides with a significant rise in temperature. A rise in Twould
reduce the strength of the rocks and allow buoyancy forces to bring
the rocks back to mid crustal levels. Further exhumation can then beexplained by a combination of extensional tectonics and erosion. The
change from rapid to slow exhumation rates described in the
section on geochronology may correspond to such a change in exhu-mation mechanism. Several authors have suggested that the forma-
tion of the Sanbagawa metamorphism is related to the approach and
subduction of an active spreading ridge (Aoya, Uehara, Matsumoto,Wallis, & Enami, 2003). However, new dating shows that the age ofthe slab at the time of formation was around 60 Ma (Nozaki et al.,
2013) and not very young as should be the case for a ridge subduction
model.
2.20.6 |Mikabu unit
The southern boundary of the Sanbagawa terrane is marked by a
prominent series of mafic and ultramafic rocks with lesser amounts ofsiliceous mudstone, chert, and limestone collectively known as theMikabu unit. The structural relationship with the Sanbagawa terrane
has been disputed. However, recent studies suggest the Mikabu unit
is part of the overlying carapace to the Sanbagawa metamorphic ter-rane (Endo & Wallis, 2017) but has undergone low-grade Sanbagawa
metamorphism. Therefore, it is lithologically and tectonically distinct
but in terms of metamorphism can be thought of as part of the San-bagawa terrane. The Mikabu unit and similar units in the Sorachi –
Yezo terrane of Hokkaido probably represent fragments of a Jurassic
ocean plateau (Ichiyama, Ishiwatari, Kimura, Senda, & Miyamoto,2014). Igneous zircon from the Mikabu unit yields ages of ca
160–150 Ma (Endo & Yokoyama, 2019; Sawada et al., 2019).
2.21 |Shimanto terrane (including cretaceous
(K) and Paleogene to Neogene parts (P –N) (Sh)
The Shimanto terrane is a Cretaceous to Neogene accretionary com-
plex developed throughout the outer arc of southwest Japan from16 of 31 WALLIS ET AL.
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areas around Tokyo, through the Kii Peninsula, Shikoku, Kyushu, and
out to the Ryukyu Islands and has become established as one of thekey areas for studying shallow accretionary tectonics. The main rock
types are interbedded mudstone and sandstone that commonly show
well-developed sedimentary structures indicating a turbidite origin.Chert and basalt, locally including well-preserved pillow basalt, are also
locally important. The Sorachi –Yezo, Idonnappu, and Hidaka accretion-
ary complexes in Hokkaido are of a similar age to the Shimanto terraneand probably represent part of the same large plate convergent margin
(G. Kimura et al., 2016). The Sanbagawa high P/Tterrane has similar
depositional age to the Shimanto terrane and may represent the meta-morphosed equivalent (Aoki, Maruyama, Isozaki, Otoh, & Yanai, 2011;Isozaki et al., 2010; Kiminami, Hamasaki, & Matsuura, 1999).
Two major subdivisions are recognized within the Shimanto ter-
rane: an Early to Late Cretaceous northern unit and a mainly Eoceneto early Miocene southern unit. The boundary between these units is
a major tectonic boundary represented by the Aki Tectonic Line in
Shikoku, the Gobo –Hagi Tectonic Line in the Kii Peninsula, and the
Nobeoka Thrust in Kyushu. The amount of feldspar is greater in sand-
stone of the older northern unit than in sandstone of the younger
southern unit, suggesting the younger sandstone contains detritalmaterial reworked from the older one (Taira, Okada, Whitaker, &
Smith, 1982). Oceanic rocks such as chert and basalt are much more
common in the younger unit than in the older unit (Taira, Katto,Tashiro, Okamura, & Kodama, 1988).
Zones of broken formations or mélange units can be traced over
many tens of kilometers. These zones consist of a matrix of sheared
mudstone containing blocks of sandstone, chert, and basalt and arethought to represent fossil major fault zones including the ancient
plate boundary. Numerous important finds have been reported from
these rocks including the presence of pseudotachylite (Ikesawa,Sakaguchi, & Kimura, 2003; Kitamura et al., 2005; Ujiie, Yamaguchi,
Sakaguchi, & Toh, 2007), which demonstrates the presence of seismic
faulting even at relatively shallow levels in the accretionary complex.Veins at low angle to the shear plane have also been highlighted asimportant in the formation of slow earthquakes (Ujiie et al., 2018).
2.21.1 |Deformation and relationship to plate
tectonics
Folds and thrusts of the Shimanto terrane show a southward vergence
and downward younging (Kanmera & Sakai, 1975; Taira et al., 1980,
1988). The higher temperature parts of the terrane show clear north−south stretching such as in the northern unit of Kyushu. These fea-
tures are all compatible with orogen-normal tectonic movement, but
significant strike-slip movements are also possible. Large-scalenorthward-plunging antiforms form a prominent set of structures inthe younger part of the terrane. Some workers have suggested this
mega-kink structure formed by recent strike-parallel shortening
(Kano, Kosaka, Murata, & Yanai, 1990; Sugiyama, 1994; Yanai, 1986).However, G. Kimura et al. (2016) suggest they are an original feature
of the terrane associated with the subduction of sea mounts andother ocean bottom morphological features. The Shimanto terrane is
locally associated with magmatic intrusions with a similar age to thatof sedimentation; in particular, the gabbro sequence at Cape Muroto
and the granite body as Cape Ashizuri in Shikoku. G. Kimura, Hashi-
moto, Kitamura, Yamaguchi, and Koge (2014) suggest that there iswidespread intrusion at depth and this has contributed to the crustal
thickness of the region.
There are distinct gaps in the age of accreted material in the
Shimanto terrane. These may be due to tectonic erosion –a process
that has been highlighted in the history of Japan by Isozaki et al.
(2010) –but could also be explained by periods of non-accretion or
non-subduction. G. Kimura et al. (2014) propose that subductionceased between ~ 12 Ma and 8 Ma due to the change in the underly-
ing plate: from the Pacific Plate to the Philippine Sea Plate and is
responsible for one of the recognized accretion gaps.
The modern offshore region from much of the Shimanto terrane
consists of the Nankai accretionary complex, formed by the conver-
gence of the Philippine Sea Plate with the Japanese margin. Theboundary between the two is thought to be gradual but is generally
obscured by overlying sediments of forearc basins. Philippine Sea
Plate subduction beneath southwest Japan is restricted to the Neo-gene, and older parts of the Shimanto terrane are thought to have
been accreted during convergence of the Pacific and other plates
that formerly bordered southwest Japan. Former plate motion vec-tors can be reconstructed from magnetic anomalies and traces oftransform faults (Engerbretson et al., 1985), however, such methods
cannot place any good constrains on the location of plate boundaries
for sections that have been subducted away. An alternativeapproach is to compare the kinematics and age of on-land deforma-
tion with changes in plate-motion vector. This information has been
used in southwest Japan to identify the plate responsible for accre-tion of different parts of the Shimanto and Sanbagawa terranes
(Tokiwa, 2009; S. R. Wallis et al., 2009). The most likely sequence of
events is Izanagi plate ( ≥~85 Ma), Kula plate (85 –65 Ma), and Pacific
plate (65 –20 Ma) associated with sinistral, dextral, and normal con-
vergence, respectively.
The two boundaries between these three plates were extensional
and ridge subduction may be involved in various geological processes.It has been suggested that the thermal structure of the Shimanto ter-
rane represents a late stage thermal event, possibly related to ridge
subduction. However, direct evidence is scarce and there is a need fora clear set of geological criteria to determine if ridge subduction has
occurred or not.
2.22 |South Kitakami terrane (SK) including the
Matsugadaira-Motai metamorphic unit (mm) andHayachine –Miyamori opholite (hy)
The South Kitakami terrane is a large domain of dominantly continen-
tal Paleozoic rocks located in northeast Honshu and is an importantsource of information concerning Japan's geological links to the east
Asian continental margin. Although the names are similar, the NorthWALLIS ET AL. 17 of 31
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and South Kitakami terranes are geologically distinct having formed at
different times in different tectonic settings (the South Kitakami ter-rane formed as an island arc with continental basement, whereas the
North Kitakami terrane represents an accretionary complex with an
oceanic plate stratigraphy) and have contrasting lithologicalassemblages.
The main part of the South Kitakami terrane is a forearc basin
sequence with a basement of granitic rocks unconformably overlainby Silurian sediments. The Hikami Granite body is a well-known part
of the granite basement and is thought to have an age of around
440 Ma by many workers (Asakawa, Maruyama, & Yamamoto, 1999;Watanabe, Fanning, Uruno, & Kano, 1995). This age is compatiblewith the presence of a Silurian unconformity, but other workers have
suggested a younger age (M. Adachi, Suzuki, Yogo, & Yoshida, 1994;
Shimojo, Ohto, Yanai, Hirata, & Maruyama, 2010; K. Suzuki & Adachi,1991; K. Suzuki, Adachi, Sango, & Chiba, 1992), and this remains a sig-
nificant point of debate.
The sedimentary sequence consists of a basal arkose overlain
by Silurian limestone and calcareous mudstone followed by Devo-
nian interlayered mud and sandstone. Numerous fossils have been
reported from this sequence including corals, trilobites (Kato et al.,1980; T. Kobayashi & Hamada, 1977; Minato, Hunahashi,
Watanabe, & Kato, 1979), and Middle Devonian radiolarians
(Umeda, 1996). Younger Carboniferous and Permian shallowmarine strata consist of volcaniclastics, limestone, and siliciclasticdeposits with a fauna including corals and brachiopods. Higher
units yield plant fossils belonging to the Cathaysian flora (Asama,
1956, 1967) of the South China Craton.
A back arc setting in the Early Carboniferous is suggested by vol-
canism with bimodal SiO
2contents (M. Kawamura et al., 1990). The
overlying Permian sequence consists mainly of limestone and shallowmarine clastic sediments. The lack of volcaniclastic deposits contrasts
with the underlying Devonian and Carboniferous sequences. A dis-
tinctive Permian conglomerate facies consist of a muddy or sandymatrix with clasts of granite that yield ages of 260 –240 Ma
(Takeuchi & Suzuki, 2000). The presence of these granitic clasts is evi-
dence for continued arc formation and related subduction during the
Late Permian and perhaps until the Early Triassic.
Triassic to lowest Cretaceous strata in the South Kitakami ter-
rane were deposited in a shallow mar ine or alluvial environment and
a r em a i n l yc o m p o s e do fc l a s t i cr o c k si na s s o c i a t i o nw i t hr a r el i m e -stone and tuff. These strata are the youngest that are not recog-
nized in the neighboring North Kitakami terrane and give an older
limit on the age of juxtaposition of the North and South Kitakamiterranes.
2.22.1 |The Matsugataira –Motai metamorphic
unit (mm) and Hayachine –Miyamori ophiolite (hy)
A series of high P/Ttype metamorphic rocks are present in the west-
ern margin of the South Kitakami Terrane and are divided into the
Matsugadaira metamorphic rocks in the south (geographically part ofthe eastern Abukuma massif) and the Motai metamorphic rocks in the
north (geographically part of the western Kitakami massif). The meta-morphism is associated with the development of alkali-amphiboles
and pumpellyite typical for high P/T conditions (Kanisawa, 1964;
Maekawa, 1988). The main rock types represent a metamorphosedoceanic plate stratigraphy: mafic, siliceous, pelitic, and psammitic
schists, associated with metagabbro and serpentinite. The mafic rocks
have a MORB composition (Kawabe, Sugisaki, & Tanaka, 1979;Tanaka, 1975). This lithological assemblage is associated with mélange
zones suggesting an accretionary complex origin (Ehiro & Okami,
1991; Umemura & Hara, 1985). Radiometric dating of probablyrelated arc-type intrusive rocks (Isozaki et al., 2015) indicates thatsubduction in this region was active in the Cambrian. The
Matsugataira metamorphic unit is unconformably overlain by
unmetamorphosed Devonian siliciclastic rocks (Ehiro & Okami, 1990).
Two ophiolitic units lie to the east of the Motai metamorphic
rocks: the Miyamori and Hayachine Complexes (Ishiwatari et al.,
2016). The Miyamori Complex consists mainly of harzburgite anddunite, whereas the Hayachine Complex consists mainly of lherzolite.
Both complexes show unusually common development of igneous
amphibole. The Hayachine Complex also locally shows a sheeted dykecomplex with a variable mafic composition. Geochemical studies sug-
gest a forearc setting for the Miyamori Complex and a back arc setting
for the Hayachine Complex (Ishiwatari et al., 2016) implying theophiolite formed in an island arc setting (Ozawa, 2016; Uchino &Kawamura, 2016). These basement units are stratigraphically overlain
by a variety of deposits of including tuffaceous sand and mudstone
conglomerate, arkose, and limestone. The limestone is associated withSilurian corals and trilobites (Okami, Ehiro, & Oishi, 1986). Zircon
U–Pb ages of ~ 470 –450 Ma indicate an Ordovician age for the
Hayachine Complex (Shimojo et al., 2010) consistent with the age ofthe overlying sediments.
The arrangement of the Matsugataira –Motai high P/Tmetamor-
phic rocks in the west, adjacent to the forearc Miyamori Complexfollowed by the backarc Hayachine Complex further east, suggeststhe former subduction was from west to east. Structures in the neigh-
boring Nedamo accretionary terrane suggest Devonian subduction
was to the southwest suggesting a switch in subduction polarity,probably in the Silurian.
2.23 |Sorachi –Yezo terrane (SY) including the
Kamuikotan unit (k) and Horokonai ophiolite
The Sorachi –Yezo terrane is part of a large east Asian Cretaceous
forearc system that stretches ⁓1500 km from Sakhalin Island to off-
shore northeast Honshu (Ando, 2003) and in Hokkaido this terrane isoriented north −south bisecting the island. The Sorachi –Yezo terrane
consists of a Late Jurassic to Early Cretaceous ophiolite to forearc
basin sequence, which includes the Horokanai ophiolite and the over-
lying sedimentary units of the Sorachi Group and the overlying Creta-ceous forearc basin sediments of the Yezo Group. These units are
underlain by accretionary complexes that have undergone high P/T18 of 31 WALLIS ET AL.
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metamorphism and are referred to as the Kamuikotan unit
(Kiminami & Kontani, 1983), and here we treat these four units ascomponents of the Sorachi –Yezo terrane. A distinct set of Cretaceous
accretionary units and an associated ophiolite lies between the
Sorachi –Yezo and Hidaka terranes. These units may be equivalent to
the Kamuikotan unit (Iwasaki et al., 1995; Watanabe et al., 1994), but
here we discuss them as a distinct Idonnappu terrane.
2.23.1 |Horokanai ophiolite
The Horokanai ophiolite (Asahin a & Komatsu, 1979; Ishizuka, 1985) has
undergone strong tectonic disruption, but an ophiolite succession of peri-
dotite, gabbro, amphibolite, MORB-li ke basalt, and siliceous sedimentary
rocks can be defined (Ueda, 2016). This s uccession is conformably overlain
by clastic rocks of the Yezo Group. The base of the peridotite is in fault
contact with schists of the Kamuikotan unit. Basalt and gabbro underwent
low-pressure metamorphism from zeol i t et og r a n u l i t ef a c i e s .C h e r to v e r l y -
ing the basalt yields Late Jurassic radiolarians (Ishizuka, Okamura, & Saito,
1983; Kawabata, 1988; Kiminami, Komatsu, et al., 1992).
2.23.2 |Sorachi Group
The Sorachi Group is generally divided into a basaltic lower part and a
volcaniclastic upper part (Kito et al., 1986). The lower part consists
almost entirely of pillow and massive basalt and their volcaniclastic
equivalents. These units can be correlated with the basaltic section ofthe Horokanai ophiolite. The basalts are characterized by non-vesicular
and aphyric tholeiites, and are locally picritic. The origin is disputed
and ocean floor (Ishizuka, 1981; Kiminami, Kito, et al. 1985; Kiminami,Kontani, et al., 1985; Niida, 1992), oceanic plateau (Ichiyama et al.,
2012, 2014; Nagahashi & Miyashita, 2002), and back-arc (Takashima,
Nishi, & Yoshida, 2002) settings have all been proposed.
The upper part of the Sorachi Group is characterized by siliceous
and volcanogenic sedimentary rocks such as chert, siliceous mud-
stone, felsic tuff, and volcanic sandstone and conglomerate, with
locally intercalated basaltic pillow lavas, dolerite sills (Kito, 1987;Niida, 1992), felsic tuff breccia, and quartz diorite dykes (Girard,
Jolivet, Nakagawa, Aguirre, & Niida, 1991). Siliceous sedimentary
rocks yield Late Jurassic to middle Early Cretaceous radiolarians(Kanie, Taketani, Sakai, & Miyata, 1981; Kiminami, Kito, et al., 1985;
Kiminami, Kontani, et al., 1985; Kito, 1987, 1995, 1997; Minoura,
Kumano, Kito, Kamata, & Kato, 1982; Takashima, Yoshida, & Nishi,2001). The geochemistry of volcanic rocks of the upper Sorachi Group
suggests formation in an arc setting (Girard et al., 1991; Niida, 1992;
Takashima et al., 2002).
2.23.3 |Yezo Group
The Yezo Group is characterized by thick terrigeneous clastic
sequences ranging from Early Cretaceous to Palaeocene. It is wellknown for its rich fossil contents –in particular ammonoids and
inoceramids –and well-preserved sedimentary and paleoenvironmental
records (Ando, 2003; Takashima et al., 2004). The lower part of the
Yezo Group conformably overlies the Sorachi Group, and its base is
defined as the onset of terrigenous clastic supply represented bysiliciclastic sandstone or black mudstone (Kanie et al., 1981; Kito,
1987). These stratigraphic relationships are the main basis for assigning
the Yezo Group a forearc basin origin (Kiminami, Kito, et al., 1985;Kiminami, Kontani, et al., 1985; Niida & Kito, 1986). The lower Creta-
ceous sediments consist of un-fossiliferous deepsea turbidities followed
by turbidities with calcareous micro- and mega-fossils. The upper Cre-taceous units are represented by strongly bioturbated shelf mudstoneand sandstone with abundant molluscan fossils. The uppermost parts
are sandy and gravelly, reflecting shallower water deposition. Sedimen-
tary blocks of serpentinite (H. Nagata, Kito, & Nakagawa, 1987;K. Yoshida et al., 2010; K. Yoshida, Taki, Iba, Sugawara, & Hidaka,
2003), blueschist, and epidote amphibolite (M. Kawamura, Ueda, &
Narushima, 1999; Ueda, Kawamura, & Yoshida, 2002) are locallyobserved in coarse sedimentary layers implying syn-orogenic exhuma-
tion of mantle material and high P/Trocks.
2.23.4 |Kamuikotan unit
The Kamuikotan unit consists of high P/Tmetamorphic rocks and ser-
pentinite, exposed in the cores of anticlines within the Yezo Group. Most
of the high P/Trocks occur either as coherent schist domains associated
with lower-grade metamorphic accr etionary units, or as blocks within
serpentinite mélange (Sakakibara & Ota, 1994). The Kamuikotan meta-
morphic rocks can be grouped into at least three types in terms of prot-
olith, metamorphic condition, and ages (Ueda, 2016).
1. The first group represents the oldest rocks that consist of epidote-
albite amphibolite with minor garnet amphibolite and garnet-quartz schist and occur as isolated blocks in serpentinite mélangeor in lower-grade schists (Bikerman, Minato, & Hunahashi, 1971;
Ishizuka & Imaizumi, 1980; Katoh, Niida, & Watanabe, 1979;
Nakagawa & Toda, 1987). Mica Ar cooling ages of 130 –125 Ma
(Bikerman et al., 1971; Imaizumi & Ueda, 1981; Ota, Sakakibara, &
Itaya, 1993) likely record a time after peak metamorphic tempera-
ture. These lithologies commonly show retrograde blueschist faciesmetamorphism –for example, sodic amphibole rimming horn-
blende –implying an anticlockwise P–Tpath, as expected shortly
after the onset of subduction.
2. The second group consists dominantly of metabasites and chert
associated with pelitic and psammitic rocks and occurs mainly as
coherent masses such as the Iwashimizu Complex (M. Kawamuraet al., 1998; Ueda, 2005). These rocks underwent blueschist faciesand lower grade (lawsonite-albite facies) metamorphism character-
ized by the presence of metamorphic aragonite, lawsonite, sodic
amphibole, and jadeitic pyroxene. Chert of the Iwashimizu Com-plex yields Late Triassic earliest Cretaceous radiolarians. Mudstone
yields early to mid-Cretaceous radiolarians and this is likely toWALLIS ET AL. 19 of 31
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represent the age of accretion (Hori & Sakakibara, 1994;
M. Kawamura et al., 2001). Metabasites of the Iwashimizu Com-plex show oceanic island geochemical characteristics (G. Kimura,
Sakakibara, & Okamura, 1994; Nakano & Komatsu, 1979;
Sakakibara et al., 1999). Similarities in sedimentary ages and basaltgeochemistry suggest links between the Iwashimizu Complex and
the accretionary units in the Idonnappu terrane (Iwasaki et al.,
1995; A. Okamoto et al., 2015; Watanabe et al., 1994).
3. The third group is found in the Pankehoronai Complex and is char-
acterized by strongly foliated and folded pelitic schist with lesser
amounts of similarly deformed siliceous, psammitic, and maficschists. This group is characterized by pumpellyite-actinolite faciesmetamorphism. A higher thermal gradient than the Iwashimizu
Complex is indicated by the presence of calcite instead of aragonite
as the stable form of CaCO
3and actinolite instead of glaucophane.
2.24 |Tokoro terrane (Tk)
The Tokoro terrane lies to the east of the Hidaka terrane in Hokkaido
and can be separated into the Nikoro Group which is an accretionarycomplex and the Saroma Group which is the overlying forearc basin
deposit (Kiminami, Kito, et al., 1985; Kiminami & Kontani, 1983;
Kiminami, Kontani, et al., 1985). The Nikoro Group consists mainly ofmetabasalt with chert, local pelagic limestone, and volcaniclastic rocks.The pelagic sediments yield Late Jurassic to Early Cretaceous micro
fossils (Iwata et al., 1990; Kiminami, Suizu, & Kontani, 1983; Okada
et al., 1989) and the clastic deposits are likely to be Upper Cretaceous(Sakakibara, Isozaki, Nanayama, & Narui, 1993). The mafic rocks of
the Nikoro Group show a mixture of tholeiitic and alkaline basaltic
compositions, which may represent abyssal ocean floor and seamountorigins, respectively (Niida, 1981; Sakakibara et al., 1986), or ocean
floor affected by plume-ridge interaction (Yamasaki & Nanayama,
2017). Zeolite to prehnite-pumpellyite facies metamorphism is wide-spread and associated with the development of metamorphic arago-nite, lawsonite, sodic pyroxene, and sodic amphibole. The estimated
metamorphic conditions are 200 –300
/C14C at 0.5 –0.6 GPa (Sakakibara,
1991). The Saroma Group is a coherent sedimentary sequence of con-glomerate, sandstone, and mudstone and occurs in synclines uncon-
formably overlying the Nikoro Group. A Late Cretaceous sedimentary
age is suggested by bivalve (Obata, Hayami, Matsukawa, Teraoka, &Taketani, 1993) and radiolarian fossils (Kanamatsu, Nanayama,
Iwata, & Fujiwara, 1992).
2.25 |Ultra-Tanba terrane (UT)
The Ultra-Tanba terrane is a Permian accretionary complex that
occurs sandwiched between the Maizuru and Mino –Tanba –Ashio ter-
ranes. It is mainly distributed in central Honshu, but can be correlated
with the older parts of the North Kitakami –Oshima terrane (Nakae &
Kurihara, 2011) and the Sikhote –Alin Mountains, eastern Russia
(Kojima, Kemkin, Kametaka, & Ando, 2000). The oldest units of theUltra-Tanba terrane are coral-bearing Carboniferous limestone associ-
ated with radiolarian chert and basalt (Igi, Kuroda, & Hattori, 1961;Pillai & Ishiga, 1987; Takemura, Suzuki, & Ishiga, 1993). Chert of Perm-
ian age is also recognized (Sugamori, 2009) as associated with fossilifer-
ous middle to late Permian mudstone sandstone and felsic tuff.
Three distinct tectonic units can be defined separated by thrust
faults. The sedimentary age of chert and mudstone shows an overall
younging from higher to lower structural units compatible with a sub-duction accretion history. The northward dip of the slaty cleavage and
geometries of mesoscopic folds suggest a southward transport direc-
tion. Finite strain and degree of metamorphism increase downsection (K. Kimura, 1988; Takemura & Suzuki, 1996). The relationshipbetween the Ultra-Tanba and Mino –Tanba –Ashio terrane is not well
understood. A recent finding in the Ultra-Tanba terrane of conglomer-
ate with serpentinite pebbles probably derived from the Oeyamaophiolite suggests that the Early Paleozoic ophiolite was exposed in
the forearc of the Permian subduction zone in which the Ultra-Tanba
terrane was accreted, while the Permian Yakuno ophiolite formed atthe same time in the volcanic arc –back arc areas (Sugamori &
Ishiwatari, 2015).
3|ZONAL ARRANGEMENT OF JAPAN
Although there are complications, probably due to post-accretion tec-
tonics, the zonal arrangement of the Japanese terranes as described
above generally follows an oceanward-younging rule, which corre-
sponds to downward younging in cross section. This arrangementreflects successive underplating of accreted terranes beneath the
hanging wall of the subduction zone –a process that operated
throughout the Phanerozoic. As an example, in the Inner Zone ofsouthwest Japan, from north to south, the early Paleozoic Oeyama
ophiolite is tectonically underlain by the metamorphic rocks of the
Sangun terrane, which is in turn underlain by the late PaleozoicYakuno ophiolite and Permian accretionary complexes of the Akiyoshiand Ultra-Tanba terranes. The Jurassic accretionary complexes of the
Mino –Tanba –Ashio terrane occupy the lowermost position. In the
Outer Zone of southwest Japan, this pattern is thought to be repeatedby the Kurosegawa and Chichibu terranes, which are further underlain
by the Cretaceous-Neogene accretionary complexes of the Shimanto
terrane. Analogous zonal arrangements are observed in northeastJapan (the Tohoku –Hokkaido area) and are also reported from other
areas of the circum-Pacific orogenic belts such as the western United
States and northeastern Russia.
ACKNOWLEDGMENTS
The authors are very grateful to T. Uchino and T. Nishiyama whosethoughtful and detailed reviews greatly helped to improve the clarityof this contribution. The authors are also grateful to Y. Iryu and
T. Muto for their editorial guidance. Finally, the authors acknowledge
the contributors to chapters 1, 2(a) –(g), and 3 of the book the Geology
of Japan –the information presented by these authors formed the
basis for many parts of this summary.20 of 31 WALLIS ET AL.
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ORCID
Simon R. Wallis https://orcid.org/0000-0002-7875-5740
Ken Yamaoka https://orcid.org/0000-0002-3609-9087
Hayato Ueda https://orcid.org/0000-0002-4267-0737
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How to cite this article: Wallis SR, Yamaoka K, Mori H,
Ishiwatari A, Miyazaki K, Ueda H. The basement geology ofJapan from A to Z. Island Arc . 2020;29:e12339.
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Gondwana Research 74 (2019) 90-100 Contents lists available at ScienceDirect GRNDEARCA Gondwana Research ELSEVIER journal homepage: www.elsevier.com/locate/gr Tectonic setting required for the preservation of sedimentary mélanges in Palaeozoic and Mesozoic accretionary complexes of southwest Japan Koji Wakita Yamaguchi University,Japan ARTICLE INFO ABSTRACT Article history: The Palaeozoic to Mesozoic accretionary complexes of southwest Japan include various types of melange. Most Received 19 August 2018 melanges are polygenetic in origin, being sedimentary or diapiric melanges that were overprinted by tectonic de- Received in revised form 12 March 2019 formation during subduction. Sedimentary melanges, without a tectonic overprint, are present in the Permian ac- Accepted 17 March 2019 cretionary complexes of the Akiyoshi and Kurosegawa belts and in the Early Cretaceous accretionary complex of Available online 25 March 2019 the Chichibu Belt. These melanges are characterized by dominant basalt and limestone clasts, within a mudstone matrix. The basalt and limestone clasts within the sedimentary melanges were derived from ancient seamounts. Subduction of a seamount results in deformation of the pre-existing accretionary wedge, and it is difficult to in- corporate a seamount into an accretionary wedge; therefore, preservation of seamount fragments requires a spe- cial tectonic setting. Oceanic plateau accretion might play an important role in interupting the processes of subduction and accretion during the formation of accretionary complexes. Especially the Mikabu oceanic plateau might have caused the cessation of accretion during the Early Cretaceous. The subduction and accretion of volca- nic arcs and oceanic plateaux helps to preserve sedimentary melanges from tectonic overprinting by preventing further subduction. ? 2019 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. 1. Introduction complexes. However, in rare cases in ancient accretionary complexes of Japan, sedimentary melanges have been preserved without a tectonic Mélanges are complex geological units composed of a mixture of overprint, including Permian and Early Cretaceous accretionary com- various rock types. The origin of melanges has been discussed by various plexes of Southwest Japan. authors (e.g., Raymond, 1984; Cowan, 1985; Festa et al., 2010, 2012; Diapiric melange is another type of chaotic body in the ancient ac- Codegone et al., 2012: Dilek et al.,2012; Wakabayashi, 2012). Sedimen- cretionary complexes (e.g., Barber et al., 1986; Barber and Brown, tary processes are one of the major causes of melange formation 1988). Diapiric melange, without a tectonic overprint, is difficult to dis- (e.g, 0kamura, 1991; Pini, 2004; Camerlenghi and Pini, 2009; Ogata tinguish from sedimentary melange. However, the former includes et al., 2012). After the formation of chaotic mixture by the sedimentary clasts, into which mudstone matrix is forcefully injected because of ex- processes how chaotic structures may be preserved. Franciscan melange traordinary pore pressure. Therefore, we can recognize the diapiric me- or complex is one of the most famous chaotic units in the world. Re- lange by the observation of the shape of melange clasts as well as the cently various processes were proposed for the chaotic sedimentary occurrence of melange within the surrounding rock units. In this mixtures of the Franciscan complex (Wakabayashi, 2011, 2012; paper, the authors choose sedimentary melange as the research target, Raymond and Bero, 2015; Platt, 2014; Ernst, 2016; Krohe, 2017). after careful observation of textures, structures and type of Submarine sliding is common along convergent margins worldwide, emplacement. and generates muddy chaotic deposits in trenches or on trench slopes This paper describes the nature of sedimentary melanges in (e.g., McAdoo et al., 2000). Ancient accretionary complexes include sed- Palaeozoic to Mesozoic accretionary complexes of Japan and discusses imentary melanges as well as other types of melange. During the devel- the tectonic setting required to preserve sedimentary melanges without opment of accretionary wedges, such chaotic units, after having been tectonic overprinting, based on an analysis of limestone and basalt formed by sedimentary processes, are affected by tectonic shearing blocks within melanges. along various types of fault. Tectonic overprinting leads to difficulty in identifying the origin of such chaotic units in ancient accretionary 2. Melanges in Japanese accretionary complexes The Japanese Islands are composed mainly of ancient accretionary E-mail address: k-wakita@yamaguchi-u.ac.jp. complexes of various ages (Isozaki et al., 2010a, 2010b, 2011; Wakita, https:/doi.org/10.1016/j-gr.2019.03.006 1342-937X/@ 2019 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. K.Wakita/Gondwana Research 74(2019) 90-100 91 2013, 2018; Wakita et al., 2018a; Fig. 1). Mélanges are common components of these accretionary complexes. Two types of melange Stage6 Cenozoic occur in the Japanese Islands: those with a mudstone matrix and those with a serpentinite matrix. The former were formed mainly 66.0 Stage5 by fragmentation and mixing of “ocean plate stratigraphy" (e.g., Isozaki et al., 1990; Matsuda and Isozaki, 1991; Wakita and Cretaceous Metcalfe, 2005; Kusky et al., 2013), and are called ocean plate stratig- raphy (OPS) mélanges (Wakita, 2015). The main components of OPS 145.0 Stage4b mélanges are basalt, limestone, chert, siliceous shale, sandstone and mudstone, which originally formed the OPs. The OPS was scraped off Jurassic along a decollement and accreted onto an accretionary wedge, and 201.3 Stage4a part of the OPS was fragmented and chaotically mixed to form OPS Triassic mélanges (Wakita, 2013, 2015). The accretionary process is suggested to be one of the main formation mechanisms of OPS 251.9 melanges, which are recognized in accretionary complexes in several Stage3 Permian tectonic belts in Japan, such as the Nedamo, Akiyoshi, Kurosegawa, Mino-Tamba, Chichibu and Shimanto belts (Fig. 1). Serpentinite 298.9 melanges, which are the second type of melange, consist of tectonic Carboniferous mixtures of various rocks and formations. Various authors have Stage2 suggested that tectonic erosion of pre-existing accretionary 358.9 complexes along the convergent margin was the main cause of Fig. 2. Stage division of Japanese accretionary complexes and stratigraphic positions of formation of serpentinite melanges in Japan (Tsujimori and Itaya, sedimentary melanges without tectonic overprinting. Red stars show the accretion age 1999; Is0zaki et al., 2010a, 2010b, 2011). of Akiyoshi Limestone and Funafuseyama Limestone in the accretionary complexes. The formation of accretionary complexes of the Japanese Islands is Stars show the accretionary age of atoll carbonates (limestone). divided into six stages on the basis of age and tectonics (Fig. 2; Wakita et al., 2018a). Accretionary complexes of Stage 1 occur in the Kurosegawa and Hida-Gaien belts in Early Palaeozoic time and are after melange formation. Accretionary complexes of Stage 2 are recog- strongly metamorphosed. The internal structures of the melanges in nized in the Nedamo and Nagato-Renge belts. The melanges of the these belts are highly deformed and the belts were metamorphosed Nedamo Belt comprise chert and basalt in a mudstone matrix (Uchino et al., 2005; Uchino and Kawamura, 2010). The Permian accretionary complexes of Stage 3 are characterized by huge limestone blocks derived from atoll carbonates on top of volcanic seamounts in the 135°E 145°E Panthalassa Ocean (Sano and Kanmera, 1988). Mélanges containing Cenozoic volcanic arc 45°N large blocks of atoll carbonate are found in the Permian accretionary NeogenetoQuaternary complexes of the Akiyoshi and Kurosegawa belts. Stage 4 is subdivided accretionarycomplex into two substages: Stage 4a and Stage 4b. The Jurassic and Early Creta- middle Cretaceous to Paleogene ceous accretionary complexes of Stage 4 include various types of accretionary complex mélange (e.g., Wakita, 2012) that contain clasts and blocks of sandstone, M.Triassic to early Cretaceous chert, limestone and basalt of varying size in a fine-grained matrix accretionary complex (Wakita, 2012). The melanges were originally of sedimentary, diapiric Permian- Early Triassic and/or tectonic origin, but most were tectonically deformed after their accretionary complex 40°N formation. The Late Cretaceous to Palaeogene accretionary complex of Permian arc Stage 5 is part of the Shimanto Belt, which includes other renowned melanges (e.g., Kimura and Mukai, 1991). The melanges of the L.Devonian to Carboniferous Shimanto Belt are mainly tectonic melanges that resulted from plate subduction along the decollement, and have been moved upwards Paleozoic arc along out-of-sequence thrusts. The accretionary complexes of Stage 6 Continentalfragment do not contain melanges. 35°N Fig.5 3. Melanges containing atoll carbonate rocks 35°N 3.1. Melanges of the Shingai Formation in the Kurosegawa Belt Fig.9 The melanges of the Shingai Formation in the Kurosegawa Belt Figs.3,4,7 are typical sedimentary melanges and contain large blocks of limestone and basalt derived from accreted seamounts (Figs. 3 and 0100200km 30°N 4; Is0zaki, 1986; Wakita et al., 2007; Hara et al., 2018). It was previously assigned to the Northern Chichibu Belt (Isozaki, 1986), 135°E1 140°EI but in this paper is regarded as part of the Kurosegawa Belt. The Shingai unit includes variously sized clasts and blocks of limestone, Fig. 1. Distribution of major accretionary complexes in Japan. OE: Oeyama Ophiolite, KB: limestone breccia, basalt, chert, sandstone, siliceous mudstone, Kurosegawa Belt, ND: Nedamo Belt, NR: Renge Metamorphic Belt, AK: Akiyoshi Belt, SO: schist and granite in a mudstone matrix (Isozaki, 1986). The matrix Suo Metamorphic Belt, Mz: Maizuru Belt, UT: Ultra Tamba Belt, MT: Mino-Tamba Belt, is not fissile and is weakly deformed. The clasts and blocks range in RK: Ryoke Belt, CB: Chichibu Belt, SB1: Sambagawa Metamorphic Belt, SB2: Shimanto Metamorphic Belt, SM: Shimanto Belt. size from a few millimetres to several kilometres. The large blocks (Modified from Wakita, 2013) are mainly basalt and limestone (Fig. 4). K.Wakita/GondwanaResearch74(2019)90-100 -33°40'N Kam Shirakidani N Shogase NangokuA Sambosan Kochi Konan 10km 133°30'E Tosa ShimantoBelt Kashiwagi unit chert Sambagawa Belt N.Chichibu Belt limestone Mikabu Belt S.Chichibu Belt basalt 133°30'E Sambosan unit PermianAcc. KurosegawaB. ncluding Shingai Formatic Fig. 3. Geological map of the south Shikoku region, showing the Shingai Formation of the Kurosegawa Belt and the Sambosan unit of the Chichibu Belt. (Modified from Hara et al., 2018) 130°30E 13535'E 33°36'N Koubu River N Kochi Station mudstone Kagami River Kochi City chert limestone 4km basalt 33°32'N Fig. 4. Geological map of the Shingai Formation. (Modified from Geological Survey of Japan AIST, 2018) K.Wakita/GondwanaResearch74(2019)90-100 93 131°10E 131°20E N 34°15'N Akiyoshidai Tsunemqri sedimentary melange Tsunemori Formation Mine 34°10'N sandstone chert 4km limestone basalt Fig. 5. Geological map of the Tsunemori Formation. (Modified from Geological Survey of Japan AIST, 2018) 3.2.Melanges of theTsunemoriFormation of limestone, limestone breccia, basalt, chert, sandstone, siliceous mud- stone, schist and granite within a mudstone matrix (Fig. 7; Isozaki, Small outcrops of melange in the Tsunemori Formation are interca- 1986). The mudstone matrix is usually not fissile and is less deformed lated with mudstone-dominated turbidite sequences (Figs. 5 and 6). than the matrix of older mélanges in the Chichibu Belt (Fig. 8). The clasts The melanges contain centimetre-sized pebbles of sandstone and lime- and blocks range in size from a few millimetres to several kilometres. stone in a mudstone matrix. It is generally only weakly deformed and is Basalt and limestone are the main mappable blocks in the melange. not fissile (Wakita et al., 2018b). 3.4.Melanges of the Funafuseyama unit,Mino-Tamba Belt 3.3. Melanges of the Sambosan unit of the Chichibu Belt The Jurassic accretionary complex of the Mino-Tamba Belt includes The Sambosan unit is the youngest unit of the Jurassic to Early Creta- various types of mélange (Wakita, 2000). These melanges are divided ceous accretionary complex of the Chichibu Belt (Matsuoka et al., 1998; into two major types: sandstone-chert mélanges and limestone- Wakita et al., 2007). The unit includes variously sized clasts and blocks basalt mélanges. The former is of diapiric origin (Wakita, 1988a, 2cm Fig. 6. (left) Sedimentary melange in the Tsunemori Formation in outcrop and (right) pebbly mudstone of the matrix in hand specimen. K.Wakita/Gondwana Research74(2019)90-100 133°28E 133°28E 133°29E 33°30'30” Ten-no Otani 33°30'N Hatta melange matrix sandstone EarlyCretaceous chert sedimentary melange of the limestone Sambosan unit Hiro-okaKami Hiro-okaNaka basalt 500 1000m Jurassicaccretionarycomplex Fig. 7. Geological map of the Sambosan unit, southern Chichibu Belt. (Modified from Geological Survey of Japan AIST, 2018) 1988b), the latter of tectonic origin (Fukui and Kano, 2006). The Some large limestone bodies directly overlie basalt in a single tectonic limestone-basalt melanges of the Funafuseyama unit of the Mino- block or slab in the accretionary complexes. Therefore, the nature of Tamba Belt are characterized by large blocks of limestone and basalt the basalt provides important information on melange emplacement. with minor amounts of chert and sandstone in a black mudstone matrix Ogawa and Taniguchi (1989) analysed the geochemistry of basaltic (Figs. 9 and 10; Wakita, 1991). The matrix is black in colour and perva- rocks in Japanese accretionary complexes of various ages. They demon- sively sheared, with a high carbon content. The melange includes very strated that most of the basalt of the Akiyoshi, Mino-Tamba and few terrigenous fragments such as quartz and feldspar. The pelitic ma- Chichibu belts is alkaline basalt of hot-spot origin. However, Jones trix is deformed and shows a scaly fabric dominated by penetrative et al. (1993） analysed rare earth elements in basalt of the Mino Belt shear fractures, with some rotated clasts. Quartz veins cut some of the and concluded that the tectonic setting of the basalt within the Jurassic shear planes and are displaced by later-stage shear planes. Most of the accretionary complex of the Mino Belt was a spreading-axis-centred mappable blocks of limestone and basalt were tectonically detached Oceanic plateau or ridge, similar to Iceland or the Ninety-East Ridge. during accretion, rather than during submarine sliding. Koizumi and Ishiwatari (2006) suggested that the Type I Suite of the Tamba Belt contains basalts with enriched mid-ocean ridge basalt (E- 4. Basaltic rocks in Japanese accretionary complexes MORB), ocean island tholeite (OIT) and ocean island alkali (OIA) basalt affinities, and that the Type II suite of the Tamba Belt is composed of OIT Basaltic rocks are key to understanding the tectonic setting of mé- and OIA basalt. They proposed the oceanic plateau subduction model to langes that contain atoll carbonate rocks. Most of these melanges yield explain the mixing of basalt and limestone in the Type II Suite of the blocks of basalt, together with limestone, as the main exotic blocks. Tamba Belt. Ichiyama et al. (2008, 2014) supported the idea that Fig. 8. Outcrop of the sedimentary melange of the Sambosan unit, Chichibu Belt. Research74(2019)90-100 36.5° rassicaccretionarycomplex Quaternary sandstone nelangematrix Fig. 9. Geological map of the Funafuseyama unit, Mino-Tamba Belt. subduction of an oceanic plateau contributed to the formation of the Ju- on the Izanagi Plate. Ichiyama and Ishiwatari (2004) suggested that rassic accretionary complex of the Mino-Tamba Belt. Koizumi and the Yakuno Ophiolite is of back-arc origin, based on the geochemistry Ishiwatari (2006) hypothesized that a seamount or oceanic island of of mafic rocks of the ophiolite. However, Herzig et al. (1997) and Suda hot-spot origin was formed either on, or adjacent to, the oceanic pla- (2004) reported Permian plagiogranite from the Yakuno Ophiolite, indi- teau, and that slices of the thicker plateau were emplaced as tectonic cating the occurrence of intra-oceanic volcanic arc activity during Perm- slabs in the melanges of the accretionary complex. ian time. Koide et al. (1987) suggested that the Maizuru Belt represents Although the Mikabu Ophiolite and Yakuno Ophiolite do not occur an east-west transect through an oceanic island at an oceanic ridge, an as tectonic blocks in melanges, they might be related to melange forma- oceanic ridge, an island arc and a marginal sea during the Permian. tion in coeval accretionary complexes. The Mikabu Ophiolite is the metamorphic counterpart of the Sambosan unit of the Chichibu Belt 5. Lithosphere and seamount ages, and seamount height (Maruyama et al., 1997; Isozaki et al., 2010a, 2010b). Accretion of mafic and ultramafic rocks occurred during the Early Cretaceous, con- It is difficult to determine the age of an oceanic plate below a sea- temporaneous with the formation of the Sambosan unit melanges. mount and the age of eruption of the seamount. Few dates are available Ozawa et al. (1997） and Ozawa (1999) suggested that basalt of the for basalt in seamounts incorporated into Japanese accretionary com- Mikabu Ophiolite originated from normal MORB or an oceanic plateau, plexes; however, it is possible to estimate the age of seamount forma- and Ichiyama et al. (2014) regarded the Mikabu Belt as a plume-type tion and the age of the oceanic lithosphere where the seamount ophiolite that was derived from a Late Jurassic oceanic plateau formed erupted by using the age range and thickness of limestone developed a Fig. 10. Outcrop of melange (a) and pillow lava (b) of the Funafuseyama unit, Mino-Tamba Belt K.Wakita/Gondwana Research 74(2019) 90-100 Permian accretionary complex (Akiyoshi Belt) Jurassic accretionary complex(Mino Belt) Age of accretion : late Middle Permian 260Ma Age of accretion:MiddleJurassic17oMa (Time range ca. 90Myr) (Time range ca. 20Myr) (Thickness > 1000m) (Thickness > 800m) late Middle Permian 260Ma - -260MalateMiddlePermian early Early Carboniferous 350Ma- 280Malate Early Permian Fig. 11. Comparison of the age range and thicknesses of limestone in Permian and Jurassic accretionary complexes. on the seamounts. Oceanic lithosphere becomes gradually cooler, seamount to accretion. As there is no major hiatus in deposition and thicker and heavier after its formation. As the oceanic lithosphere the limestone formed continuously, the top of the seamount remained sinks gradually because of its weight, the ocean overlying the litho- near the surface of the ocean. Therefore, the age range and thickness sphere becomes gradually deeper. Persons and Sclater (1977) and of the Akiyoshi Limestone indicate that the oceanic plate subsided by n sin s is s o io ui oo age and ocean depth. McKenzie et al. (2005) supported the former part of the GDH-1 curve in Fig. 12. Based on this fitting, the oceanic model, whereas Miller et al. (2008a, 2008b) supported the latter. This plate was formed ~120 Myr prior to its subduction. The seamount was paper uses the GDH-1 model (Global Depth and Heat flow model) pro- formed 30 Myr after the formation of the oceanic plate. Based on this es- posed by Stein and Stein (1992). GDH-1 model is more useful for prac- timation, the oceanic plate on the seamount originated at ~380 Ma (Late tical calculation on depth and age of oceanic plate. Devonian). Although long-term sea-level fluctuations should also be In the Permian accretionary complex of the Akiyoshi Belt and the Ju- considered (Miller et al., 2008), this method is useful for estimating rassic accretionary complex of the Mino Belt, limestone generally occurs the age of the oceanic plate and the timing of seamount eruption. as blocks or slabs. The present study examined the largest slab of lime- Large limestone blocks overlying basaltic rock are also present in the stone in each accretionary complex. The Akiyoshi Limestone of the Jurassic accretionary complex of the Mino Belt. The basaltic rocks of the Akiyoshi Belt is about 1000 m thick and ranges in age from the earliest Jurassic accretionary complexes occur as tectonic blocks or slabs within Carboniferous to the late middle Permian (Fig. 11; Nakazawa and the melange matrix, together with limestone and chert blocks in the Ueno, 2004). The Akiyoshi Limestone is underlain by basalt, and the Funafuseyama unit of the Mino Belt. The limestone blocks in the mé- age of the youngest part of the limestone is close to the age of the mud- langes of the Funafuseyama unit are Permian in age, ranging from stone that forms the matrix of the mélange. The mudstone was trench Sakmarian (Pseudoschwagerina zone） to Capitanian (Fig. 11; Yabeina sediment when the limestone was accreted to the continental margin. zone; Wakita, 1988b; Sano, 1988). The age of the basaltic rocks can be Thus, the limestone records the entire history from the birth of the estimated from the age of the overlying limestone. The oldest ages of (A) (B) 3000- 3000- a1200m (height of seamount) ca 2000m (height of seamount) 三 one formation) 4000- 4000- 300m(thickness of limestone) estone formation) 10Ma(age length bet 5000 1000m 5000 thickness of limestone) Oce mount formation 6000- -0009 Cessation of limestoneformatior Arrival totr rench Arrival to trench 20 40 60 180 20 40 60 80 100120140160180 0 Oceanic crustal age [Myr] Oceanic crustal age [Myr] Fig. 12. Estimates of the ages of seamounts and oceanic plates that formed Permian and Jurassic accretionary complexes in Japan (after McKenzie et al., 2005). (a) Akiyoshi Belt, (b) Mino Belt. The curved line is the GDH-1 model of Stein and Stein (1992), and shows the age-depth relationship for an oceanic plate. Limestone thicknesses and ages are illustrated in Fig. 11. K.Wakita/GondwanaResearch74(2019)90-100 97 the limestone blocks must be close to the age of the limestone associ- landslide in the Hawaian Islands described by Moore (1964). This pro- ated with the basalt. This means the basaltic rocks of the Middle Jurassic cess is the first possible mode of fragmentation of atoll carbonate before mélanges of the Mino Belt must be latest Carboniferous to earliest Perm- its accretion (Fig. 13). ian in age. 15.2 Lallemand et al. (1989) showed that the Daichi-Kashima Seamount and seamount eruption in the Jurassic accretionary complex. The collapsed at the trench, because of bending of the oceanic plate. Normal Funafuseyama Limestone of the Mino Belt is one of the largest and best- faults have developed on the surface of the Pacific Plate near the Japan studied limestone blocks in the Jurassic accretionary complex. The block Trench, and the Daichi-Kashima Seamount on the Pacific Plate has is about 800 m thick and ranges in age from late early Permian to late been cut by normal faults and collapsed at the Japan Trench. Wakita middle Permian (Sano, 1988). The age range and thickness of the (1988a, 1988b) proposed a diapiric process for mélange formation in Funafuseyama Limestone indicate that the oceanic plate subsided by the Jurassic accretionary complex of the Mino Belt. However, in his re- ~800 m over 20 Myr. This subsidence rate is consistent with the grey vised model, fragmentation of both atoll carbonate and seamount basalt part of the GDH-1 curve in Fig. 12. Based on this fitting, the oceanic occurred in the deeper part of the accretionary wedge, along the plate was formed about 130 Myr prior to its subduction. The seamount decollement of the tectonic blocks of limestone and basalt in the Jurassic was formed 10 Myr after the formation of the oceanic plate. This result accretionary complex of the Mino-Tamba Belt (Wakita, 2000, 2012, is consistent with the conclusion of Jones et al. (1993) that the basalt of 2015). The structural analysis of Kimura and Mukai (1991) indicated the Mino Belt originated as a bathymetrically elevated feature located at that fragmentation of the ocean plate stratigraphy occurred along the or near a mid-oceanic spreading axis. Koizumi and Ishiwatari (2006) sug- gested that an oceanic plateau originated by either within-plate Palaeogene accretionary complex of the Shimanto Belt. Although basalt, superplume magmatism or along the plume-influenced segment of the chert and sandstone are tectonically fragmented, few blocks of lime- palaeo-Pacific spreading ridge. Based on this estimation, the oceanic stone are recognized in the mélanges. plate on the seamount was formed at ~290 Ma (Early Devonian). This limestone reached the sea surface and stopped growing in the late middle 6.2. Mixing Permian, and moved continuously towards the trench until the Middle Ju- rassic. Thus, the top of the seamount was below sea level for 110 Myr, al- s jo ansae s ns es a eu paeade d pos a yno bonate of Carboniferous to Permian age at the trench, and suggested level change. This inference is supported by the presence of mixed faunas that atoll carbonates collapsed at the trench and became mixed with of Permian and Triassic age within the uppermost parts of other limestone trench sediments to form melanges in the Permian accretionary com- blocks in the Jurassic accretionary complex (e.g., Matsuda, 1980). cess for mélange formation in the Shingai Formation of the unit of the Mino Belt contain minor alkaline basaltic rocks (Hattori Kurosegawa Belt, based on the weakly deformed nature of the mélange and Yoshimura, 1983), and Wakita (1984) reported basaltic blocks of matrix. Permian limestone and basalt were gradually mixed with other Triassic to Jurassic age that are much younger than the other major ba- components of the Jurassic accretionary complex, during development saltic blocks, which were accreted during the Jurassic. These basaltic of the accretionary wedge (Wakita, 1988a, 1988b, 2000, 2006, 2012, 2015). Mixing of chert and basalt with sandstone and mudstone oc- some of them are alkaline. curred along the decollement during the development of the Late Creta- In the Cretaceous accretionary complex of the Shimanto Belt, basaltic ceous melanges of the Shimanto Belt. The fragmentation and mixing rocks are associated with chert rather than with limestone. However, that caused the development of tectonic melanges in the Shimanto ac- both basalt and chert generally occur as independent tectonic blocks cretionary complex resulted from simultaneous shearing in the within melanges. Even the oldest age of the chert cannot be used to con- decollement zone. strain the exact age of the basalt, as chert and basalt are independent tec- tonic blocks detached from different part of the ocean plate stratigraphy. 6.3.Accretion 6. Possible setting of fragmentation, mixing and accretion of atoll carbonate to form melanges in ancient accretionary complexes. The melanges of 6.1. Fragmentation the Jurassic accretionary complex of the Mino-Tamba-Chichibu Belt were formed by this type of progressive melange formation (Wakita, After the formation of volcanic islands in the ocean, it is possible for volcanic islands with atolls to collapse into the ocean before they arrive tigraphy was gradually deformed, finally generating completely chaotic at the subduction trench, forming structures such as the giant undersea mixtures (Fig. 14). submarine sliding atoll seamount trench carbonate collapse sediment seamount collapse limestone basalt basalt pelagic sediment oceanicplate nderplating Fig. 13. Possible tectonic settings for the fragmentation and mixing of atoll carbonate 98 K.Wakita/Gondwana Research 74(2019) 90-100 melangeswith limestone and basalt decollement trench fill pelagic seamount Themipelagic Oceanic Plate Offscraping of atoll carbonate and top of seamount Fig. 14. Melange formation by underplating of seamount fragments. (After Wakita, 2012) However, it is difficult for atoll carbonate to be accreted at a trench. in mélanges of the Jurassic accretionary complex of the Mino Belt Scholl and von Huene (2007) observed that no accretion occurs in 75% (Wakita, 2012). This is common, in cases in which atoll carbonates com- of present-day convergent margins. In most modern convergent mar- gins, tectonic erosion is more common than accretion. As mentioned stone and basalt blocks. above, fragmentation and mixing of atoll carbonate on the Daichi- However, a special setting is required for the formation and preser- Kashima Seamount with trench sediments occurs at the Japan Trench, vation of sedimentary melanges without a tectonic overprint. The Perm- where tectonic erosion occurs rather than accretion. In contrast, the ian mélanges of the Akiyoshi and Kurosegawa belts and the Early Nankai Trough in southwest Japan is a site of sediment accretion. At Cretaceous melanges of the Chichibu Belt include huge limestone and the Nankai Trough, seamounts collide with soft sediments of the basalt blocks within a mudstone matrix, that exhibits neither shearing young accretionary wedge and are subducted into the deeper part of nor tectonic overprinting. This evidence suggests that the limestone the accretionary wedge, together with the slab, without any deforma- and basalt blocks were detached and accreted near the trench or in tion (Yamazaki and Okamura, 1991). In both examples of convergent the shallow part of the accretionary wedge, and that they did not suffer margins, at the Japan Trench and the Nankai Trough, atoll carbonate tectonic shearing caused by progressive plate subduction. that detached from seamounts cannot enter the accretionary wedge The sedimentary melanges that include limestone and basalt blocks near the trench. In order for atoll carbonate to be accreted at a trench, developed at the end of the formation of the accretionary complexes in it is necessary to consider a special tectonic setting in which atoll car- each stage (Fig. 2). The sedimentary mélanges of the Shingai Formation bonates are fragmented and mixed with sediments at the trench, but and of the Tsunemori Formation were formed during the late Permian tectonic erosion does not occur. In the next section, a possible tectonic in the Kurosegawa Belt and the Akiyoshi Belt, respectively. These sedi- setting for sedimentary mixing and accretion of atoll carbonates with- mentary melanges were formed during the latest stage in the accretion- out tectonic erosion is discussed. ary process. The melanges of the Sambosan unit contain Early Cretaceous fossils 7. Possible setting for the formation of sedimentary melanges with- in their matrix, which indicates that the melanges are the youngest out tectonic overprinting components in the Chichibu Belt. The formation of sedimentary mé- langes that include limestone and basalt blocks also occurred in the lat- Formation of a sedimentary melange, without tectonic overprinting, est stage of the accretionary process during the Early Jurassic to Early might take place as follows. A carbonate atoll develops on a volcanic Cretaceous (Fig. 2). seamount in the ocean. The volcanic seamount and the atoll carbonate The Japanese Islands are composed of accretionary complexes of move towards the trench, together with the oceanic plate. In the case various ages. The major components are Permian, Late Triassic to Early of the Hawaiian-Emperor Chain, atoll development ceases when the Cretaceous, and Late Cretaceous to Palaeogene accretionary complexes. volcanic island subsides beneath sea level (Grigg, 1997). However, the There are several reasons for disruptions in the formation of accretion- ary complexes, such as ocean ridge subduction, tectonic erosion and the early Carboniferous to middle Permian (Nakazawa et al., 2009). volcanic activity (Isozaki et al., 2010a; Wakita et al., 2018a). To account The direction and speed of plate movement in the late Palaeozoic for the presence of sedimentary melanges without a tectonic overprint, Panthalassa Ocean might have been different from present-day plate it is necessary to envisage a special tectonic setting in Japan during the movement. late Permian and Early Cretaceous. After a long period of travel on the oceanic plate, a carbonate atoll on Flat-slab subduction might be the key to understanding the preser- top of a volcanic seamount arrives at a trench. Some of the atoll carbon- vation of sedimentary mélanges containing atoll carbonates. Flat-slab ate might collapse at the trench due to normal faulting in the upper sur- subduction was proposed to account for the Late Jurassic to Early Creta- face of the oceanic plate. However, these seamount collapses occur at ceous igneous record of East Asia (Kiminami and Imaoka, 2013; Kim the tectonic erosion boundary in modern oceanic settings (Lallemand et al., 2016), although the reason for proposing flat subduction is un- et al., 1989). clear. A buoyant mass is necessary to cause flat subduction. A candidate At accretionary convergent boundaries, seamounts are subducted for the buoyant mass in the case of the Japanese Islands in Early Creta- without deformation, and atoll carbonate can be detached from the sea- ceous is the Mikabu oceanic plateau (Maruyama et al., 1997; mount by decollement in the deeper part of the accretionary wedge Utsunomiya et al., 2008; Safonova et al., 2009). The Mikabu Belt is char- (Fig. 14). Such atoll carbonates have been recorded as tectonic blocks acterized by igneous rock of basaltic composition. The basic igneous K.Wakita/GondwanaResearch74(2019)90-100 99 Oceanic Plateau seamounts References Barber, AJ, Brown, K, 1988. Mud diapirism: the origin of melanges in accretionary com- plex? Geology Today 4, 89-94. Barber, AJ, Tjokrosapoetro, S., Charton, TR, 1986. Mud Volcanoes, Shale Diapirs, Wrench Faults, and Melanges in Accretionary Complexes, Eastern Indonesia. AAPG Bulletin 70, 1729-1741. Camerlenghi, A, Pini, AA, 2009. Mud volcanoes, olistostromes and Argill scagliose in the Mediterranean region. Sedimentology 56, 319-365. Codegone, G., Festa, A., Dilek, Y., Pini, G.A., 2012. Small-scale polygenetic melanges in the pirism in superposed melange evolution in orogenic belts. Tectonophysics 568, Oceanic Plateau Collision 170-184. Cowan, D.s., 1985.Structural styles in Mesozoic and Cenozoic melanges in the western Cordillera of North America. Geological Society of America Bulletin 96, 451-462. Preservation of Dilek, Y, Festa, A, Ogawa, Y, Pini, G.A., 2012. Chaos and geodynamics: melanges, mé- Sedimentary Melange lange-forming processes and the significance in the geological record. Tectonophysics 568-569, 1-6. SubductionJum Ernst, W.G, 2016.Franciscan melanges: coherent blocks in a low-density, ductile matrix. International Geology Review58， 626-642.H https://doi.org/10.1080/ 0026814.2015.1108879. Festa, A., Pini, G.A., Dilek, Y., Codegone, G., 2010. Melanges and melange-forming pro- cesses: a historical overview and new concepts. International Geological Review 52, 1040-1105. Festa, A., Dilek, Y., Pini, G.A., Codegone, G., Ogata, K., 2012. Mechanisms and process of a e s o p p n tions: Redefining and classfying melanges. Tectonophysics 568, 7-24. Fig. 15. Possible scenario for the preservation of sedimentary mélanges that include Fukui, ,KanK,6Defmation rocs andkinatis fmlange inthEaly seamount fragments. tonics 26, 1-14. Geological Survey of Japan AIST, 2018. Seamless Geologic Map of Japan. V2 https:/gbank. gsjjp/seamless/. Grigg, R.W., 1997. Paleoceanography of coral reefs in the Hawaiian-Emperor Chain - rocks of the Mikabu Belt are considered to be a fragment of an ancient revisited. Coral Reefs 16, S33-S38. oceanic plateau that formed during the Jurassic and accreted during the Early Cretaceous (Maruyama et al., 1997; Isozaki et al., 2010a). logical Map f Japan, 1:00o0Kochi (2nd ediion).Glogical Surveyof Japan,IT The subduction and accretion of the Mikabu oceanic plateau occurred, Hattori, I., Yoshimura, M, 1983. Late Triassic to Middle Jurassic ages for greenstones within the MesozoicNanjo Massif of theMino terrane,Central Japan.Memoirf Fac together with the formation and preservation of sedimentary melanges, ulty of Education, Fukui University, Series II, no. 32, 67-80 (in Japanese with English that do not exhibit tectonic overprinting (Fig. 15). abstract). To preserve sedimentary melanges, including limestone and basalt Herzig, C.T, Kimbrough, D.L., Hayasaka, Y, 1997. Early Permian zircon uranium-lead ages for plagiogranites in the Yakuno ophiolite, Asago district, Southwest Japan. Island Arc blocks, it is necessary that oceanic plate subduction is halted. If oceanic 6, 396-403. plate subduction had continued, even after oceanic plateau collision, the Ichiyama, Y., Ishiwatari, A., 2004. Petrochemical evidence for off-ridge magmatism in a back-arc setting from the Yakuno ophioite, Japan. The Island Arc 13, 157-177. sedimentary melanges would have been affected by tectonic deforma- Ichiyama, Y., Ishiwatari,A, Koizumi, K,008.Petrogenesis of greenstones from the Min- tion due to accretion and the development of thrust fault systems in Tamba belt, SW Japan: evidence for an accreted Permian oceanic plateau. Lithos 100, the accretionary prism. Therefore, the presence of sedimentary mé- 127-146. langes without tectonic shearing suggests that the site of oceanic plate Ichiyama, Y., Ishiwatari, A, Kimura,J, Snda, R, Miyamoto, T,2014, Jrassic plume-origin ophiolites inJapan: accretedfragments ofoceanplateau, Contributions toMineralogy subduction jumped oceanwards because of the buoyancy of the incor- porated oceanic plateau (Fig. 15). Isozaki, Y., 1986. The Shingai Formation in the northern Chichibu Belt, Southwest Japan, The oceanic plateaux responsible for preservation of the Permian Landmass. Journal of Geological Society of Japan 92, 497-516 (In Japanese with En- sedimentary mélanges of the Akiyoshi and Kurosegawa belts are un- glish abstract). known. However, an oceanic plateau or other buoyant mass must Isozaki, Y., Maruyama, S., Fukuoka, F, 1990. Accreted ocean materials in Japan. have been subducted and incorporated into the accretionary wedges Tectonophysics 181, 179-205. Isozaki, Y., Aoki, K., Nakama, T., Yanai, S., 2010a. New insight into a subduction-related during the late Permian or Early Triassic time. A Permian volcanic arc orogen: a reappraisal of the geotectonic framework and evolution of the Japanese complex existed in the Maizuru Belt, which is located near the Permian Islands. Gondwana Research 18, 82-105. accretionary complex of the Akiyoshi Belt (Fig. 1). The Maizuru volcanic Isozaki Y., Maruyama, S., Aoki, K, Nakama, T, Miyashita, A, Otoh, S, 2010b.Geotectonic arc might provide clues to the cause of flat subduction during the late ments and boundaries of Pacific-type (Miyashiro-type) Orogen. 119, 999-1053. Permian. Isozaki,,Muyama ,Nakama,Ymmt,anai,211rwthand hink The subduction and collision of oceanic plateaux or island arcs of an active continental margin: updated geotectonic history of the Japanese Islands. Journal of Geography (Chigaku Zasshi) 120, 65-99. changed the conditions of plate subduction, causing a cessation or re- Jones, G San,H,aamiJn , 9aure and ttnic etting f acreed basa duction in volcanic activity along the ancient convergent margin of from the Mino terrane, Central Japan. Journal of the Geological Society 150, Proto-Japan. Thus, sedimentary melanges without a tectonic overprint 1167-1181. in ancient accretionary complexes represent a marked alteration in Kim, S.W., Kwon, S., Park, S-, Le, C., Cho,D.L, Lee, H-J, Ko, K., Kim, SJ, 2016. SHRIM U-Pb dating and geochemistry of the Cretaceous plutonic rocks, Korean Peninsula: a the plate tectonic setting of Japan. new tectonic model of the Cretaceous Korean Peninsula. Lithos 262, 88-106. Kiminami, K., Imaoka, T, 2013. Spatiotemporal variations of Jurassic-Cretaceous Acknowledgements magmatism in eastern Asia (Tan-Lu Fault to SW Japan): evidence for flat-slab sub- duction and slab rollback. Terra Nova 25, 414-422. Kimura, G., Mukai, A, 1991. Underplated units in an accretionary complex: melange of I thank Dr. Anthony J. Barber of Royal Holloway, University of London, UK and Dr. Yujiro Ogawa of Tsukuba University, Japan for their Koide, Y, Tazaki, K, Kagami, H, 1987. Sr isotopic study of Ibara dismembered ophiolite from the Maizuru Tectonic elt, Southwest Japan. Journal of the Japanese Association critical reviews of this paper. I also wish to acknowledge Drs. Andrea of Mineralogists, Petrologists and Economic Geologists 82, 1-15. Festa of the University of Torino, Italy, Kei Ogata of the University Centre Koizumi, K, Ishiwatari, A, 2006. Oceanic plateau accretion inferred from Late Paleozoic in Svalbard, Norway, Gian Andrea Pini of the University of Trieste, Italy greenstones in the Jurassic Tamba accretionary complex, Southwest Japan. Island and Yildirim Dilek of Miami University, USA for providing me the oppor- Arc 15, 58-83. Krohe, A., 2017. The Franciscan Complex (California, USA) - the model case for tunity to submit this paper. This work was supported by JSPS KAKENHI return-flow in a subduction channel put to the test. Gondwana Research 45, Grant Number JP16K05579. 282-307. 100 K.Wakita/Gondwana Research 74(2019)90-100 Kusky, TM., Windley, B.F,Safonova, I., Wakita, K.,J, Polat, A,Santosh, M, 2013.Recogni Sano, H., Kanmera, K, 1988. Paleogeographic reconstruction of accreted oceanic rocks, tion of ocean plate stratigraphy in accretionary orogens through Earth history: a re- Akiyoshi, Southwest Japan. Geology 16, 600-603. cord of 3.8 billion years of sea floor spreading, subduction and accretion. Sano,H, Kanmera,K, 1991.Collapse f ancient ceanic reef comlex -what happened Gondwana Research, 24, 501-547. during collision of Akiyoshi reef complex? - sequence of collisional collapse and gen- Lallemand, S., Culotta, R, von Hune, R, 1989. Subduction of the Daichi Kashima Sea- eration of collapse products. 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Tectonics of an Early Carboniferous forearc inferred from McKenzie, D., Jackson, J., Priestley, K., 2005. Thermal structure of ocean and continental a high-P/T schist-bearing conglomerate in the Nedamo Terrane, Northeast Japan. Is- lithosphere. Earth and Planetary Science Letters 233, 335-349. land Arc 19, 177-191. Moore, JG., 1964. Giant submarine landslides on the Hawaian Ridge. US Geological Sur- Uchino, T., Kurihara, T., Kawamura, M, 2005. Early Carboniferous radiolarians discovered vey Professional Paper 501-D, D95-D98. from the Hayachine Terrane, Northeast Japan: the oldest fosil age for clasticrocks of Miller, RD., Sdrolias, M, Gaina, C., Roest, W.R, 2008a. Geochemistry. Geophysics, accretionary complex in Japan. Journal of Geological Society of Japan 111, 249-252 (In Japanese with English abstract). Miller, RD., Srolias, M., Gaina, C., Stinbrger, B.,Heine, C, 200b.Long-tem seavl Utsunomiya, A, Suzuki, N., Ota, T, 2008. 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Wakita 2019 Tectonic setting required for the preservation of sedimentary melanges in Paleozoic and Mesozoic accretionary complexes of southwest Japan.txt
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Research Article
Blueschist-facies metamorphism during Paleozoic orogeny in
southwestern Japan: Phengite K–Ar ages of blueschist-facies tectonic
blocks in a serpentinite melange beneath early Paleozoic Oeyama
ophiolite
TATSUKI TSUJIMORI1AND TETSUMARU ITAYA2
1Department of Earth Sciences, Faculty of Sciences, Kanazawa University, Kanazawa 920–1192 and
2Research Institute of Natural Science, Okayama University of Science, Okayama 700–0005, Japan
Abstract Blueschist-bearing Osayama serpentinite melange develops beneath a peridotite
body of the Oeyama ophiolite which occupies the highest position structurally in the centralChugoku Mountains. The blueschist-facies tectonic blocks within the serpentinite melangeare divided into the lawsonite–pumpellyite grade, lower epidote grade and higher epidotegrade by the mineral assemblages of basic schists. The higher epidote-grade block is agarnet–glaucophane schist including eclogite-facies relic minerals and retrogressive law-sonite–pumpellyite-grade minerals. Gabbroic blocks derived from the Oeyama ophioliteare also enclosed as tectonic blocks in the serpentinite matrix and have experienced ablueschist metamorphism together with the other blueschist blocks. The mineralogic andparagenetic features of the Osayama blueschists are compatible with a hypothesis thatthey were derived from a coherent blueschist-facies metamorphic sequence, formed in asubduction zone with a low geothermal gradient ( ~10°C/km). Phengite K–Ar ages of 16
pelitic and one basic schists yield 289–327 Ma and concentrate around 320 Ma regardlessof protolith and metamorphic grade, suggesting quick exhumation of the schists at ca
320 Ma. These petrologic and geochronologic features suggest that the Osayamablueschists comprise a low-grade portion of the Carboniferous Renge metamorphic belt. The Osayama blueschists indicate that the ‘cold’ subduction type (Franciscan type)metamorphism to reach eclogite-facies and subsequent quick exhumation took place in the northwestern Pacific margin in Carboniferous time, like some other circum-Pacific oro-genic belts (western USA and eastern Australia), where such subduction metamorphismalready started as early as the Ordovician.
Key words: K–Ar phengite age, Osayama blueschist, Oeyama ophiolite, Paleozoic orogeny ,
Renge metamorphic belt, serpentinite melange.
Cotkin et al. 1992), eastern Australia ( ca 480 Ma,
Fukui et al. 1995) and Kurosegawa klippe, south-
western Japan ( ca 350–390 Ma, Ueda et al. 1980).
Paleozoic blueschists in these regions are alwaysassociated with Paleozoic ophiolite, and occur generally as tectonic blocks in a serpentinitemelange. Recent studies on Paleozoic ophiolites inthe circum-Pacific region have documented theophiolite formation in a supra-subduction zonesetting (forearc, volcanic arc, or back arc; Ozawa1988; Arai & Yurimoto 1994; Wallin & Metcalf1998). Tsujimori (1998) also showed that someINTRODUCTION
The blueschist-facies metamorphic rocks provide
critical evidence for paleo-subduction zones. In thecircum-Pacific orogenic belts, the incipient sub-duction of the paleo-Pacific plate took place duringthe Early–Middle Paleozoic, as indicated by theblueschist-facies metamorphic rocks from theKlamath Mountains, western USA ( ca 450 Ma,
Accepted for publication November 1998.
© 1999 Blackwell Science Asia Pty Ltd.The Island Arc (1999) 8,190–205
Paleozoic blueschist-facies metamorphism in SW Japan 191
ophiolitic fragments enclosed in the serpentinite
melange have experienced a blueschist metamor-phism together with other blueschist blocks. Thissuggests that the hanging wall of a subductionzone was also dragged into depths by tectonicerosion and metamorphosed.
In central Chugoku Mountains, southwestern
Japan, a serpentinite melange bearing blueschistblocks of various metamorphic grade (Osayamaserpentinite melange, Tsujimori 1998) developsbeneath the Early Paleozoic Oeyama ophiolite. This setting is a good example for studying subduction and exhumation processes through a joint geochronologic–petrologic method. Thispaper presents newly obtained K–Ar age data for
the blueschist-facies tectonic blocks from theOsayama serpentinite melange and discusses thetectonic implications of the Paleozoic orogeny insouthwestern Japan.
GEOLOGICAL SETTING
Southwestern Japan is a well-developed, circum-
Pacific type orogenic belt with oceanward growthof the accretionary complex since the middle Paleozoic (Fig. 1), as determined by a large amount of field-geological and biostratigraphic
Fig. 1 Geotectonic subdivision of Southwest Japan (modified from Isozaki & Itaya 1991; Isozaki & Maruyama 1991). Localities of the Pale ozoic
blueschists are also shown. Hd, Hida low-P/T metamorphic belt; Ok, Oki low-P/T metamorphic belt; Oe, Oeyama ophiolite; HM, Hida marginal belt; Rn,
Renge high-P/T metamorphic belt; Ak, Akiyoshi accretionary complex; Mz, Maizuru belt (Yakuno ophiolite); Ut, Ultra-Tamba accret ionary complex; So,
Suo high-P/T metamorphic belt; M-T, Mino-Tamba accretionary complex; Ry, Ryoke low-P/T metamorphic belt; Sb, Sambagawa high-P/T metamorphic
belt; Cn, northern Chichibu accretionary complex; Ks, Kurosegawa belt; Cs, southern Chichibu accretionary complex; Sh, Shimanto accretionary complex;
MTL, Median Tectonic Line; ISTL, Itoigawa–Shizuoka Tectonic Line.
14401738, 1999, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.1999.00231.x by Ohio State University University Libraries, Wiley Online Library on [17/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
data (Hayasaka 1987; Nishimura 1990; Ishiwatari
1991; Isozaki & Itaya 1991; Isozaki 1996; Nakajima1997). Paleozoic ophiolite and blueschist have beensporadically distributed in the Chugoku Moun-tains, occupying the highest structural positions inthe nappe pile.
RENGE BLUESCHIST
Geochronologic data accumulated since the 1980sled to the subdivision of the ‘Sangun metamorphicbelt’ into two or three discrete units (Watanabe et al. 1987; Hayasaka 1987; Shibata & Nishimura
1989; Nishimura 1990; Isozaki & Maruyama 1991;Nakajima 1997). Most recently , Nishimura (1998)divided the ‘Sangun metamorphic belt’ into twobelts: the Renge belt (330–280 Ma) and the Suobelt (230–160 Ma). We follow the terminology ofNishimura (1998) for the high-P/T schist belts inthe Inner Zone of southwestern Japan. However,we distinguish the associated ophiolitic peridotitebodies from the Renge and Suo belts of Nishimura(1998) as ‘Oeyama ophiolite’, because they clearlypre-date the schists and have different tectono-metamorphic history .
The Renge blueschists in the Chugoku Moun-
tains occur as thin nappes, which are overlain bythe Oeyama ophiolite, and also appear as tectonicblocks within the serpentinite melange beneaththe Oeyama nappe (Fig. 2). The sporadic outcropsof the Renge blueschists comprise a disrupted
metamorphic belt, which has been considered asthe western extension of the blueschist-bearingOmi serpentinite melange. The Renge blueschistsmay have constituted a late Paleozoic regionalhigh-P/T metamorphic belt, which has been fragmented during exhumation and nappeemplacement.
OEYAMA OPHIOLITE
The peridotite bodies of the ‘Oeyama ophiolite’occupy the structurally highest position in theChugoku Mountains (Fig. 2). They are composedmainly of moderately depleted harzburgite (residual spinel peridotite) and dunite with gabbroic intrusions (diallage gabbro and dolerite).The eastern peridotite bodies such as at Oeyamahave slightly more fertile features than thewestern bodies, such as Tari-Misaka and Osayama(Arai 1980; Kurokawa 1985; Nozaka & Shibata1994; Matsumoto et al. 1995; Tsujimori 1998). The
podiform chromitites enclosed in dunite are char-acteristically developed only in western peridotite(Tari-Misaka body , Arai 1980; Matsumoto et al.
1997), and amphibolites (metacumulate and gneissose metagabbro) occur as tectonic block only in eastern peridotite bodies (Oeyama body , Kurokawa 1985; Wakasa body , Nishimura &Shibata 1989). The Ochiai–Hokubo body in the192 T . Tsujimori and T . Itaya
Fig. 2 (a) Distribution of geotectonic
nappe pile in eastern Chugoku Moun-tains. (b) Distribution of the peridotitebodies of the Oeyama ophiolite in thecentral Chugoku Mountains. Black areasrepresent ultramafic bodies.
14401738, 1999, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.1999.00231.x by Ohio State University University Libraries, Wiley Online Library on [17/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Paleozoic blueschist-facies metamorphism in SW Japan 193
central Chugoku Mountains, which is quite differ-
ent in petrologic features from the other bodies(Arai et al. 1988; Matsukage & Arai 1997), may be
a different geological unit in view of the radiomet-ric ages of gabbroic rocks (237–245 Ma, Nishimura& Shibata 1989).
The residual peridotites are mineralogically
very similar to the estimated mantle restites ofback-arc basin basalt and mid-ocean ridge basalt(MORB; Arai 1994). The presence of high-Al pod-iform chromitites and petrologic features of theresidual peridotite strongly indicate that theOeyama ophiolite represents a supra-subductionzone ophiolite formed in the back-arc basin orprimitive arc setting (Arai & Yurimoto 1994, 1995;Matsumoto et al. 1997; Zhou et al. 1998). The
gabbroic intrusions crosscutting peridotites of the Oeyama ophiolite show a MORB-like majorelement pattern (Hayasaka et al. 1995), which isalso compatible with the back-arc basin setting of
the Oeyama ophiolite.
GEOLOGY OF THE OSAYAMA SERPENTINITE
MELANGE
The Osayama serpentinite melange develops
beneath the Osayama peridotite body (Fig. 3). Theserpentinite melange is tectonically underlain by the Suo schists, and is in contact with theunmetamorphosed, molasse-type shallow marine sediments of the Jurassic Y amaoku Formation(Konishi 1954) on the north by a high-angle fault.All these rocks are unconformably overlain by theEarly Cretaceous Kyomiyama conglomerate. Themassive peridotite unit and the Suo schists haveundergone an overprint of contact metamorphismby Cretaceous granitic intrusives on the west.
Fig. 3 Geological map of the Osayama serpentinite melange (after Tsujimori 1998). Phengite K–Ar ages obtained in this present paper ar e also shown
in parentheses. Shaded area by broken lines represents the metamorphic zones in contact aureole by Cretaceous granites after No zaka and Shibata (1995).
LP , schist of lawsonite–pumpellyite grade; E, schist of epidote grade; Gb, diallage gabbro; Dl, dolerite; At, albitite.
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The blueschist-facies schists, fragments of the
Oeyama ophiolite (serpentinized peridotite,gabbro, dolerite) and metasomatic rocks (albitite,jadeitite, omphacitite, tremolite schist etc.) areenclosed as tectonic blocks of various size (10 cm to 1.5 km in length) in serpentinite matrixconsisting of schistose, friable and fine-grainedserpentinite with pebble-to-boulder-size frag-ments of serpentinized peridotite. The peridotiteblocks contain Fo
90.5–91.5 olivine, orthopyroxene with
2.4–3.0 wt% Al 2O3and chromian spinel with a
Cr/(Cr +Al) ratio of 0.40–57 as primary minerals.
Petrologic features of melange matrix peridotitesuggest that the melange matrix has been derivedfrom widely varying western peridotite bodies ofthe Oeyama ophiolite such as the Tari-Misaka andAshidachi bodies (Tsujimori 1998). Hashimoto andIgi (1970) first described lawsonite–glaucophaneschists in the eastern part of the serpentinitemelange here studied. The blueschist-faciesschists are divided into the lawsonite–pumpellyitegrade and epidote grades based on the mineralassemblages of basic schists intercalated in the pelitic schist. They correspond to the lawsonite–blueschist and epidote–blueschist faciesvarieties of Evans (1990), respectively . The epidotegrades contain two varieties, a garnet-free lower-grade block and a garnet-bearing higher-gradeblock (garnet–glaucophane schist). The blocks ofthe lawsonite–pumpellyite grade are the mostdominant type. The gabbro and dolerite blocksalso contain blueschist-facies mineral assemblagesof lawsonite–pumpellyite grade, but the gabbroicintrusives in the neighboring peridotite body donot have any blueschist-facies high-P/T minerals.The gabbroic blocks often grade into the basic schist of lawsonite–pumpellyite grade withincreasing textural deformation. The chemistry ofthe igneous clinopyroxenes and bulk rock compo-sitions of the Osayama gabbroic blocks indicatethat the blocks have been derived from the gabbroic intrusions of the Oeyama ophiolite (Tsujimori 1998).
PETROLOGY OF BLUESCHIST-FACIES BLOCK
LAWSONITE–PUMPELLYITE GRADE (LAWSONITE–
GLAUCOPHANE SCHIST AND GABBROIC FRAGMENTS)
The lawsonite–pumpellyite grade blocks are
characterized by the assemblage Na-amphibole +lawsonite or Na-amphibole +pumpellyite in
basic schists, although the mineral assemblage andtexture are variable from block to block. The basic
schists include the following mineral assemblageswith albite, quartz and titanite in excess: Na-amphibole +lawsonite +chlorite +phengite, Na-
amphibole +lawsonite +pumpellyite +chlorite,
Na-amphibole +lawsonite +pumpellyite +stilp-
nomelane, Na-amphibole +pumpellyite and
Na-amphibole +chlorite. In the fine-grained
sample, albite and quartz are exactly identified byusing an electron-probe microanalyzer. Titanite,relic augite, K-feldspar, sulfides, zircon and apatiteoccur as accessory minerals in some blocks. Most of the Na-amphiboles in this grade are glaucophane to ferro-glaucophane, and their compositions are variable for different mineralassemblages (Fig. 4). One block contains zonedNa-amphibole having a glaucophane core and a ferro-glaucophane rim (Fig. 4). Evidence of agreenschist-facies overprint such as an actinoliticrim on Na-amphibole is not observed in this grade.
The pelitic schists of the lawsonite–pumpellyite
grade consist mainly of quartz, phengite, chloriteand albite with minor titanite and apatite. Car-bonaceous matter, lawsonite, K-feldspar, tourma-line and carbonate minerals occur also in someblocks. A penetrative schistosity (S
1) defined by
phengite and chlorite is commonly observed. Insome cases, fine-scale crenulation cleavage (S
2) is
developed and overprints a crenulated S 1fabric.
Although phengite of S 1fabric is finer ( <0.2 mm)
than that of S 2(0.3–0.5 mm), no compositional dif-
ferences are recognized.
The ophiolitic fragments (gabbro and dolerite)
derived from the Oeyama ophiolite also have theblueschist-facies mineral assemblage, similar to lawsonite–pumpellyite-grade basic schists.Igneous plagioclase is replaced by aggregates ofpumpellyite or lawsonite and albite and igneousilmenite altered to aggregates of titanite. Na-amphibole occurs in three modes: overgrowingepitaxially on relict augite and hornblende, fillingcracks of clinopyroxene, and replacing patchedamphiboles included in clinopyroxene.
LOWER EPIDOTE GRADE (EPIDOTE–GLAUCOPHANE
SCHIST)
The constituent minerals of this grade are
commonly much coarser than the lawsonite–pumpellyite grade schists. The basic blocks arecharacterized by the assemblage Na-amphibole +
epidote +chlorite. The basic schist of the lower
epidote grade includes the following mineral assem-194 T . Tsujimori and T . Itaya
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Paleozoic blueschist-facies metamorphism in SW Japan 195
blages with albite, quartz and titanite: Na-amphi-
bole +epidote +chlorite, Na-amphibole +epidote +
chlorite +stilpnomelane, Na-amphibole +epidote +
pumpellyite and Na-amphibole +winchite +epidote
+chlorite +stilpnomelane. Na-amphiboles in this
grade are ferro-glaucophane to glaucophane, butsome blocks of the lower epidote grade contain prograde-zoned Na-amphibole having a crossitecore and glaucophane rim (Fig. 4). Albite oftenoccurs as porphyroblasts (maximum length: 3.5 mm). Although an actinolitic rim on Na-amphibole is rarely found in some small blocks,greenschist-facies overprinting is not observed inthis grade.
The pelitic schists of the grade contain mainly
chlorite, quartz, albite, and phengite with smallamounts of epidote and titanite. Albite commonlyoccurs as porphyroblasts (0.5–2.0 mm in length)which include tiny quartz, phengite, chlorite,apatite and rarely Na-amphibole. Na-amphibole,graphite, carbonate and garnet (Prp
1–2Alm 23–33
Sps 41–57Grs 19–25) are rarely observed. A penetrative
schistosity defined by coarse-grained phengite(0.5–0.8 mm in length) and chlorite is developed.
HIGHER EPIDOTE-GRADE BLOCK (GARNET–
GLAUCOPHANE SCHIST WITH ECLOGITIC MINERALASSEMBLAGE)
The higher epidote grade is defined by the coexist-
ence of almandine-rich garnet +glaucophane andthe presence of an eclogite-facies mineral assem-
blage. In the garnet–glaucophane schist, two dis-tinct blueschist-facies stages can be defined basedon the texture and mineral zoning. The peak meta-morphic stage is characterized by the assemblageNa-amphibole (glaucophane core) +garnet +rutile
+epidote +quartz +K-feldspar. The epidote por-
phyroblasts (maximum length: 2 mm) sometimesinclude eclogite-facies mineral assemblage, garnet+omphacite (Jd
35–50Di52–56Ae <9) +rutile +quartz +
glaucophane, as tiny inclusions ( <0.03 mm). The
retrograde stage is characterized by the assem-blage Na-amphibole (ferro-glaucophane rim) +
chlorite +pumpellyite +titanite ± phengite, which
is equivalent to the lawsonite–pumpellyite grade.In some cases, the strongly sheared phengite-richpart is developed in the outcrop. Although compo-sitional zoning from glaucophane core to ferro-glaucophane rim is common, such zoning is notobserved in the phengite-rich part (Fig. 4). Retro-grade ferro-glaucophane often fills cracks ofgarnet. Garnet porphyroblasts (up to 3 mm indiameter) often contain tiny inclusions of rutileand quartz. Garnets in the glaucophane-rich parthave higher (Mg +Fe) and lower Mn contents than
garnets in the garnet-rich layer (Fig. 5). Thegarnets show prograde zoning where Fe and Mg increase and Mn decreases from core to rim, and compositions of garnets within epidote por-phyroblast corresponds to the rim of those in theglaucophane-rich part (Fig. 5). The distributionFig. 4 Compositional variations of Na-amphiboles from the Osayama blueschists in Miyashiro’s diagram of Fe3+/(Fe3++Al) vsFe2+/(Fe2++Mg). The
arrows show compositional zoning (R, rim; C, core).
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coefficients of Fe and Mg, K DGrt-Cpx, between garnet
and omphacite in the epidote vary from 8.3 to 15.9.This garnet–glaucophane schist is a high-gradeblock which was overprinted with the other low-grade blueschists. More detailed petrology ofthe garnet–glaucophane schist will be describedelsewhere.
BULK ROCK COMPOSITIONS OF THE OSAYAMA PELITIC
SCHISTS
The bulk rock composition of a typical law-
sonite–pumpellyite-grade pelitic schist and threelower epidote-grade schists were analyzed. Noremarkable difference between the two grades was recognized. As compared with the average ofthe Sambagawa pelitic schists (Goto et al. 1996),
the Osayama pelitic schists are characterized byhigher MgO (2.9–3.8 wt%), FeO* (5.2–8.1 wt%),P
2O5(0.16–0.40 wt%), K 2O (3.3–5.6 wt%) and
moderate CaO (0.8–1.4 wt%), MnO (0.08–0.17 wt%)and Al
2O3(14.7–19.3 wt%). The MgO/(MgO +FeO*)
mole ratio is 0.39–0.45. The A ¢value of AFM
diagram [(Al 2O3–3K 2O-Na 2O)/(Al 2O3–3K 2O-Na 2O+
FeO* +MgO)] varies from –0.15 to 0.00 and issignificantly lower than that of the Sambagawa
average (0.11). The Osayama pelitic schists arericher in mafic components than the Sambagawapelitic schists.
METAMORPHIC CONDITIONS
The Na-amphiboles in the Osayama blueschists
are characterized by a low Fe
3+/(Fe3++Al) ratio,
and are in the glaucophane and ferro-glaucophanefields except for those in some lawsonite–pumpellyite-grade blueschist, and the core compo-sition of some zoned Na-amphiboles in the lowerepidote grade (Fig. 4). Phengites in the Osayamablueschists have Si contents significantly higherthan that in the underlying Suo pelitic schists (Fig. 6). The compositions of Na-amphibole andphengites of the Osayama blueschist showcommon high-P/T features.
Although any geothermometers based on Fe–
Mg exchange reactions are not applicable for the lawsonite–pumpellyite grade of the Osayamablueschists, its approximate P–T condition can bededuced by the mineral assemblage. In the lawsonite–pumpellyite grade, glaucophane +law-
sonite and glaucophane +pumpellyite assem-
blages are observed and albite is stable. TheSchreinemakers’ net for the NCMASH (Na
2O-
CaO-MgO-Al 2O3-SiO 2-H 2O) system shows that the196 T . Tsujimori and T . Itaya
Fig. 5 Composition of garnets of garnet–glaucophane schist of the
higher epidote grade in the Mn–Fe–Mg ternary diagram. Matrix garnet:(d), glaucophane rich; ( s), garnet-rich. Eclogite garnet: ( w), inclusions
within epidote.
Fig. 6 Compositional variations in Al vsSi (p.f.u. for O =22) for phen-
gites from the Renge blueschists and the Suo schists. (a) Pelitic schist;(b) basic schist. Osayama blueschists: ( d), lawsonite–pumpellyite grade;
(u), lower epidote grade; ( s), higher epidote grade. Suo schist: ( +),
pelitic schists (underlying the Osayama melange).
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Paleozoic blueschist-facies metamorphism in SW Japan 197
solid solution, gives a minimum pressure of 1.5 GPa
at ~550°C for omphacite (X Jd=0.35–0.50) +quartz
assemblage without albite. The P–T condition ofthe final retrogression stage may correspond tothat of the lawsonite–pumpellyite grade.
In the typical high-P/T type metamorphic belts,
such as the Franciscan, Kamuikotan and NewCaledonian, their low-grade portions are charac-terized by the common assemblage of glaucophane+lawsonite or pumpellyite, and then the glauco-
phane +epidote assemblage becomes gradually
stable with increasing metamorphic grade(Y okoyama et al. 1986; Maruyama & Liou 1988;
Takayama 1988; Shibakusa 1989). In the Osayamablueschist blocks, the paragenetic feature of thelawsonite–pumpellyite grade and lower epidote-grade blueschists may be interpreted as a coher-ent metamorphic sequence that has undergonetypical high-P/T type metamorphism in the sub-duction zone, with a geothermal gradient close to10°C/km (Miyashiro 1994). The presence of albiteindicates that the metamorphic condition liesbelow the jadeite–quartz reaction line. Originalcoherency of the metamorphic sequence for theOsayama blueschists is also supported from thegeochronologic data described in the followingsection.
K–AR AGE DETERMINATION
The K–Ar ages were determined for 20 phengite
separates from 16 metamorphic rocks (Table 1):two basic and pelitic schists from the lawsonite–pumpellyite grade, 12 albite porphyroblast-bearing pelitic schists from the lower epidotegrade, and two phengite-rich parts of a garnet–glaucophane schist (higher epidote grade).Mineral assemblages of the rocks dated are shownin Table 1.
Rock samples were crushed with a jaw crusher
and then sieved to obtain a proper grain-size forconcentrating phengite. The sieved fraction waswashed using de-ionized water and dried in anoven at 80°C. Phengites were concentrated usingan isodynamic separator and a tapping on a paper,and the collected phengite was treated with 2 mol/L HCl to dissolve out chlorite along cleavageplanes. The acid-treated sample was then washedrepeatedly with ion-exchanged water and dried at80°C.
The K–Ar age determination was carried out at
Okayama University of Science following Nagao et al. (1984) and Itaya et al. (1991). Potassium was
Fig. 7 K content (wt%) vsK–Ar age (Ma) diagram showing the effect
of grain size and impurities in phengites. Tie-lined data are from the samesample and the shaded marks represent the coarser grained phengite sep-arate. ( s), (150/200) lawsonite–pumpellyite grade; ( u) (150/200), ( )
(100/150), lower epidote grade; ( u) (150/200), ( ) (100/150), higher
epidote grade.
glaucophane +lawsonite assemblage is stable at a
higher pressure than the pumpellyite–actinolitefacies and pumpellyite–diopside facies (Banno1998). The glaucophane +pumpellyite stability field
also appears as a subfacies in the glaucophane +
lawsonite field in NCMASH systems (Frey et al.
1991; El-Shazly 1994). The P–T condition of the law-sonite–pumpellyite grade is restricted in a field ofthe lawsonite–blueschist facies where the glauco-phane +pumpellyite +albite assemblage is stable.
In the same sense, the P–T condition of the lowerepidote grade is limited to the albite-stable field in the epidote–blueschist facies. The petrogeneticgrid proposed by Evans (1990) indicates that the lawsonite–pumpellyite grade did not reach theclosure temperature ( ~350°C) of the K–Ar phengite
system, whereas those of the lower epidote-graderocks were probably above that temperature.
In the garnet–glaucophane schist of the higher
epidote grade, assuming that the inclusions withinepidote are in equilibrium, the garnet–clinopyrox-ene Fe–Mg exchange geothermometer by Krogh(1988) gives 530–620°C at 1.3 GPa. The geobarom-eter using the breakdown of low albite to jadeite +
quartz (Ghent et al. 1987), assuming ideal Jd–Di
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198 T . Tsujimori and T . Itaya
Table 1 Mineral assemblages of the samples used for the K–Ar age determination
Meta. Grade (Facies) LP (LBS) E (lower EBS) E (upper EBS)
Lithology Basic Pelitic Albite porphyroblast-bearing pelitic schist Grt–Gln schist
Sample OS162a OS80 OS182 OS277 OS304 OS190 OS224 OS329 OS93 OS350 OS188 OS318 OS267 OS281 OS23 OS23B
Na-amphibole • — • • —————————— • •
Lawsonite • — ———————————— — —
Epidote — — •••••••••••• • •
Chlorite • • •••••••••••• • •
G a r n e t —— —— • ————————— • •Phengite • • •••••••••••• • •
Quartz • • •••••••••••• • •
Albite • • •••••••••••• • •
Titanite • • •••••••••••• • •
Rutile — — ———————————— • •
Carbonaceous matter — • — — • ———— • — • — • — —
Others Kfs Kfs
K–Ar age (Ma)
[100/150] — — 324 — — — 318 319 292 — — — — — 322 319[150/200] 315 311 308 327 324 283 273 312 300 285 315 312 315 289
LP, lawsonite–pumpellyite grade; E, epidote grade; LBS, lawsonite–blueschist facies; EBS, epidote–blueschist facies.
Phengite K–Ar ages are also represented.
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Paleozoic blueschist-facies metamorphism in SW Japan 199
analyzed by flame photometry using a 2000 ppm
Cs buffer. Argon was analyzed on a 15-cm radiussector-type mass spectrometer (HIRU) having asingle collector system with an isotopic dilutionmethod using a
38Ar spike (Itaya et al. 1991). Decay
constants for 40K to 40Ar and 40Ca, and the 40K
abundance used in age calculation are 0.581 ·
10–10/year, 4.962 ·10–10/year, and 0.0001167, respec-
tively (Steiger & Jager 1977). The results are pre-sented in Table 2 and are also shown visually inFigs 7 and 8.
Phengite K–Ar ages of the Osayama blueschist-
facies tectonic blocks of the lawsonite–pumpellyite,lower epidote and higher epidote grade yield311–315, 273–327 and 289–322 Ma, respectively ,which are concentrated around 320 Ma as a whole.The phengite separates dated have potassium con-tents ranging from 5.52 to 8.89, and most of them(16 samples) are greater than 6.5 wt% in potash. Asmentioned earlier, the samples of the lawsonite–pumpellyite grade never reached the closure tem-perature ( ~350°C) of the K–Ar phengite system,
whereas those of the epidote-grade rocks wereprobably above that temperature. Itaya and Taka-sugi (1988) argued that the K–Ar phengite age ofthe low-grade Sambagawa schists, which have not experienced a culmination temperature higherthan the closure temperature of the phengite K–Arsystem, represented the timing of exhumation/cooling ages because of the argon depletion from
phengite by ductile deformation during theexhumation of the host schists. Thus, we interpretthe K–Ar phengite ages as the exhumation/coolingage soon after the blueschist-facies metamorphism.However, the range of K–Ar ages of the phengiteseparates from the blueschist-facies schist tectonicblocks, wider than analytical error of individualanalysis. They may be due to either or both of (i) thedifferent cooling age among the schists at the timebecause of different argon depletion processesduring ductile deformation in exhumation ofschists; and (ii) the effect of impurities in the phen-gite separates because some finer grained fractionshave significantly lower potassium content andyounger age (Fig. 7). Although the phengite K–Arages of the Osayama blueschists-facies schists havesome variation, the concentration at 320 Ma indi-cates that the exhumation of schists took placeapproximately at that time.
DISCUSSION
GEOLOGICAL SIGNIFICANCE OF THE OSAYAMA
BLUESCHISTS
In southwestern Japan, late Paleozoic high-P/T
schists are sporadically distributed (Fig. 1).Table 2 Phengite K–Ar age data of the blueschist-facies tectonic blocks from the Osayama serpentinite melange
Sample no. Fraction Potassium (wt%) Rad. argon 40 (10-8cc STP/g) Age (Ma) Air cont. (%)
Lawsonite–pumpellyite grade
OS162a 150/200 6.811 –0.136 9086 –88 314.7 –6.4 1.4
OS80 150/200 6.442 –0.129 8489 –82 311.0 –6.3 1.1
Lower epidote grade (albite porphyroblast-bearing pelitic schist)
OS182 100/150 8.204 –0.164 11311 –110 324.3 –6.6 0.4
OS277 150/200 5.517 –0.110 7184 –72 307.7 –6.3 0.8
OS304 150/200 8.211 –0.164 11418 –113 326.9 –6.7 0.9
OS190 100/150 8.435 –0.169 11602 –110 323.6 –6.6 0.5
OS224 100/150 7.470 –0.149 10068 –99 317.6 –6.5 0.5
OS329 100/150 7.819 –0.156 10582 –103 318.8 –6.5 0.7
150/200 7.422 –0.148 8810 –84 282.6 –5.8 0.9
OS93 100/150 6.946 –0.139 8542 –82 292.0 –6.0 0.4
150/200 6.184 –0.124 7083 –69 273.4 –5.6 0.7
OS350 150/200 7.942 –0.159 10480 –100 311.5 –6.3 0.6
OS188 150/200 8.891 –0.178 11254 –108 299.8 –6.1 0.7
OS318 150/200 7.657 –0.153 9168 –89 284.9 –5.9 1.1
OS267 150/200 7.550 –0.151 10084 –101 315.0 –6.5 0.7
OS281 150/200 6.684 –0.314 8840 –87 312.2 –6.4 0.8
Higher epidote grade (garnet–glaucophane schist)
OS23 100/150 7.483 –0.150 10240 –99 322.1 –6.5 0.7
150/200 6.721 –0.134 8960 –89 314.5 –6.4 0.8
OS23B 100/150 7.065 –0.141 9578 –92 319.4 –6.5 0.7
150/200 5.682 –0.114 6900 –66 288.6 –5.9 0.9
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Although they generally occur as tectonic blocks
in a serpentinite melange, a series of high-P/Tschists have been considered to be the constituentsof a late Paleozoic regional high-P/T metamorphicbelt, called the Renge belt (Nishimura 1998), basedon the geochronologic data shown in Fig. 8. Thephengite K–Ar ages from the Osayama blueschistsare within the age variation of the high-P/T schistsin the Omi, Wakasa, Toyogadake, Wakamiya andKiyama areas of the Renge belt, indicating that theOsayama blueschists is a constituent of the Rengebelt.
Most of the Renge schists have recorded green-
schist, epidote–blueschist or epidote–amphibolitefacies assemblages, and the lawsonite–blueschistfacies rocks are extremely rare in the Renge belt(Banno 1958; Nishimura et al. 1983; Nakamizu et al. 1989; Nishimura 1990). The Osayama
blueschists having the assemblages glaucophane +
lawsonite or glaucophane +pumpellyite belong to
a typical high-P/T type metamorphic facies seriesformed in the subduction zone. The differences ofthe recorded P/T conditions between the Osayamablueschists (typical high-P/T type) and the otherRenge schists (intermediate high-pressure type)may be due to the following reasons: (1) local diver-sity of geothermal gradient in the subduction zone;and/or (2) the different exhumation rate over-printed various lower P/T conditions. The Rengebasic schists of the Wakasa area have barroisiterimmed by actinolite, suggesting greenschist-facies overprint after epidote–amphibolite facies(T . Tsujimori, unpubl. data), although there is no evidence of greenschist-facies overprint in the200 T . Tsujimori and T . Itaya
Fig. 8 (a) Frequency distribution of K–Ar and Rb–Sr phengite (white mica) ages of Paleozoic high-pressure schists in southwestern Japa n (compiled
from Maruyama & Ueda 1974; Maruyama et al. 1978; Shibata & Ito 1978; Ueda et al. 1980; Shibata & Nishimura 1989; Isozaki et al. 1992; Kabashima
et al. 1995; Kunugiza et al. 1997), and of K–Ar hornblende ages for the amphibolites and gabbroic intrusions of the Oeyama ophiolite (Shibata et al.
1979; Shibata 1981; Nishimura & Shibata 1989; Nishina et al. 1990). The metamorphic facies of the schists are distinguished. PA, pumpellyite–actino-
lite facies; LBS, lawsonite–blueschist facies; EBS, epidote–blueschist facies; GS/EBS, transitional facies between greenschist and epidote–blueschist
facies; EA, epidote–amphibolite facies. (b) Isotopic age relations of the Cordilleran high-pressure metamorphic belts in southw estern Japan, western USA,
and eastern Australia (compiled from Patrick & Day 1995; Isozaki & Maruyama 1991; Little et al. 1993; Fukui et al. 1995; Nishimura 1998). The time
scale is after Harland et al. (1990).
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Paleozoic blueschist-facies metamorphism in SW Japan 201
Osayama blueschists. This suggests that the
Osayama blueschists were formed in the subduc-tion zone with the lowest geothermal gradient ofthe Renge belt. Absence of the greenschist faciesoverprinting in the Osayama blueschists suggestsa quick exhumation, which is supported by thehomogeneous K–Ar age, soon after the blueschist-facies metamorphism in the subduction zone.
Recent advanced studies of the Oeyama ophio-
lite have revealed that they represent a supra-subduction zone ophiolite formed beneath theback-arc basin or primitive arc setting (Arai &Yurimoto 1994, 1995). Some fragments of theOeyama ophiolite pre-dating the Osayama blue-schists have also suffered the blueschist faciesmetamorphism together with the Osayama blue-schists as mentioned before. It follows that theOeyama ophiolitic lithosphere was close to thetrench of the Carboniferous subduction zonesystem, and was eroded and dragged down to adeeper part of the subduction zone to undergoblueschist-facies metamorphism together with theRenge schists. Such tectonic erosion of the supra-subduction zone lithosphere has been documentedin the modern subduction system of the Marianaarc–trench system (Bloomer 1983; Maekawa et al.
1995).
COMPARISON WITH OTHER PALEOZOIC HIGH-P/T
SCHISTS IN SOUTHWESTERN JAPAN
The subduction-related metamorphic rocks pre-
dating the Renge schists in southwestern Japanhave already been reported (Fig. 8). In the Kuro-segawa belt of the Outer Zone of southwesternJapan, which is interpreted as a tectonic klippeconsisting of pre-Jurassic equivalents of the InnerZone (Isozaki & Itaya 1991), the pumpellyite–glaucophane schists, epidote–barroisite schistsand epidote–hornblende schist have been reported(Maruyama & Ueda 1974; Nakajima & Maruyama1978; Nakajima et al. 1978). The former has pet-
rographic features similar to the Osayamablueschists but gives K–Ar phengite ages of352–394 Ma (Ueda et al. 1980), significantly older
than those of the Osayama schists. The latter hastwo groups of K–Ar ages; one is 317–327 Ma (Uedaet al. 1980), similar to the Renge belt; and the
other is 402–445 Ma (Maruyama & Ueda 1974)demonstrating the oldest high-pressure schists insouthwestern Japan.
Amphibolites as tectonic blocks in the Oeyama
ophiolite have undergone epidote–amphibolitefacies metamorphism and have demonstratedhornblende K–Ar ages from 469 to 336 Ma
(Kurokawa 1985; Nishimura & Shibata 1989;Nishina et al. 1990). A garnet amphibolite giving a
K–Ar biotite age of 442 Ma (Matsumoto et al. 1981)
occurs as tectonic blocks within the 320-Ma Rengeschists in Hida Mountains (Nakamizu et al. 1989).
A clinopyroxene-bearing garnet–amphibolite witha K–Ar hornblende age of 409 Ma (Y oshikura et al.
1981) occurs as tectonic blocks in the Kurosegawabelt. To reveal the timing of igneous activity of theOeyama ophiolite, the gabbroic intrusions havebeen dated. They gave the hornblende K–Ar agesfrom 343 to 239 Ma (Shibata et al. 1979; Nishina
et al. 1990). Some of those ages are clearly younger
than those of the Renge schists, namely , they contradict the fact that some fragments of theOeyama ophiolite are enclosed as tectonic blocksin the serpentinite matrix and have suffered theRenge blueschist metamorphism together with the other blueschist blocks. These young ages ofthe gabbroic intrusions are likely to be due to therejuvenation by the post-dating metamorphismbecause the igneous brown hornblende in the gabbroic intrusions is commonly rimmed by actinolite (Y amaguchi 1989). Recently , Hayasaka et al. (1995) preliminarily reported the Sm–Nd
ages of ca560 Ma for gabbroic intrusions in central
Chugoku Mountains, suggesting a time of forma-tion of the Oeyama ophiolite as the Cambrian.
BLUESCHIST-FACIES METAMORPHISM DURING
PALEOZOIC OROGENY IN SOUTHWESTERN JAPAN
In the circum-Pacific orogenic belt, Paleozoic
blueschist facies metamorphic rocks also occur inwestern USA and eastern Australia (Fig. 8). TheSkookum Gulch blueschist ( ca 450 Ma) distributed
in the Yreka Terrane, eastern Klamath Mountains,is characterized by the mineral assemblage glaucophane +lawsonite, resembling the Osayama
blueschists (Cotkin 1987). The Skookum Gulchblueschist is tectonically overlain by serpentinizedperidotite of the Cambro-Ordovician Trinity ophio-lite, and blueschist contains 570-Ma tonalite blocksderived from Trinity ophiolite (Wallin et al. 1988).
The New England Fold Belt in eastern Australiaincludes three Paleozoic subduction-related metamorphic rocks, ca 260, ca 340–310, and ca
470 Ma (Fukui et al. 1995). The oldest rocks contain
467–481-Ma epidote–glaucophane schist occurringalong the ophiolitic serpentinite melange zone inthe Glenrock–Pigna Barney area, northeasternNew South Wales (Fukui et al. 1995). They are also
closely associated with 530-Ma ophiolitic rocks
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(Aitchison et al. 1992). It is considered that an active
continental margin formed an accretionarycomplex, high-P/T schists and volcano-plutonism atthe circum-Pacific orogenic belt (Isozaki 1996;Maruyama 1997). In southwestern Japan, theOrdovician schists of the Kurosegawa belt(Maruyama & Ueda 1974) is evidence for an incipi-ent subduction of the paleo-Pacific Plate, andtypical blueschists appear from the Devonian(Ueda et al. 1980) (Fig. 8). The petrologic and
geochronologic comparison of the Paleozoic high-P/T metamorphic rocks in southwestern Japanrevealed that the geothermal gradient in the sub-duction system was relatively high to form theepidote–amphibolite facies metamorphic rocks inLate Ordovician–Silurian time (e.g. Maruyama &Ueda 1974). The low geothermal gradient to formthe blueschists in the Devonian–Carboniferouscould be attained by an active subduction of thepaleo-Pacific oceanic plate. In early history of thecircum-Pacific orogenic belt, the subduction systemin western USA and eastern Australia had reachedthe low geothermal gradient to form the blueschistsas early as the Ordovician, when the system insouthwestern Japan still had a high gradient.
The tectonic association of Paleozoic ophiolite
and Paleozoic high-P/T schist is pervasivethroughout the circum-Pacific region. This sug-gests that each orogenic belt in the circum-Pacificregion had experienced an early history similar tothat in southwestern Japan.
ACKNOWLEDGEMENTS
We express sincere thanks to Drs T . Okada and
H. Takeshita for their help in age determination at Okayama University of Science. The seniorauthor is grateful to Dr A. Ishiwatari for daily discussion and the critical reading of the manu-script. Sincere thanks are due to Prof. S. Arai for his important suggestion about peridotitebodies in the Chugoku Mountains, and to Dr K.Kunugiza for discussion about metamorphic rocksin Hida Mountains. Prof. S. Banno and Prof. Y .Nishimura are thanked for their constructive criticisms. The present study was supported financially in part by JSPS Research Fellowshipsfor young scientists.
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Island Arc - June 1999 - Tsujimori - Blueschist_facies metamorphism during Paleozoic orogeny in southwestern Japan.txt
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Gondwana Research 18(2010) 82-105 Contents lists available at ScienceDirect GRNSEAANA Gondwana Research ELSEVIER journal homepage: www.elsevier.com/locate/gr New insight into a subduction-related orogen: A reappraisal of the geotectonic framework and evolution of the Japanese Islands s s q 1 'rq o 1 *e oi onss a Department of Earth Science and Astronomy, The University of Tokyo, Komaba, Meguro,Tokyo 153-8902, Japan Department of Earth and Planetary Sciences,Tokyo Institute of Technology,Tokyo 152-8551, Japan Japan Geo-communications, Yotsuya, Tokyo 160-0004, Japan ＡRTICＬE ＩNＦＯ ABSTRACT Article history: Hokkaido, Japan. Sci. Rep. Nigata Univ., Received 22 October 2009 metamorphic processes. In Energetics of Geo- Received in revised form 20 February 2010 ting of basic volcanic rocks determined using Accepted 24 February 2010 chronology. These investigations have addressed various themes such as: 1) seismic profile of the crust Available online 10 March 2010 and mantle beneath the Japanese Islands, 2) high-precision ages of the protoliths of high-P/T metamorphic rocks, and 3) provenance of terrigenous clastics. The results have led to a number of important findings Keywords: including: 1) detection of a large mass of slab around the mantle boundary layer suggesting the long-term Subduction oceanic subduction beneath Japan, 2) confirmation of the subhorizontal piled-nappe structure for the entire Accretion Blueschist crust of Japan, 3) finding a new high-P/T metamorphosed accretionary complex unit that represents the Tectonic erosion the system cordierite-garnet-biotite. In Orogeny Hidaka metamorphic belt and its implication synthesis of these new data, this article presents a re-evaluation of the conventional geotectonic subdivision of the subduction-related orogen in Japan, re-definition of the elements and their mutual boundaries, and re- consideration of the geotectonic evolution of the Japanese Islands. In particular, the historical change in provenance suggests that proto-Japan has experienced large-scale tectonic erosion in multiple stages, and the corresponding large amounts of continental crust materials were subducted. For understanding the orogenic growth of Japan during the last ca. 500 million years, the significance of tectonic erosion coupled with continental contraction, as well as the oceanward accretionary growth, requires further attention. mental investigation of exchange equilibria in 1. Introduction Ever since Naumann (1885), the nearly 150 year-long intensive geologic mapping of the surface geology clarified that the majority of The Japanese Islands form a bow-shaped chain of islands over the Japanese crust is composed of strongly deformed sedimentary 3000 km in length along the eastern margin of Asia. Among the many rocks (dominant sandstone and mudstone with minor volcanics, modern island arc systems in the world, the Japanese Islands likely chert, and limestone), metamorphic rocks (crystalline schist, gneiss), represent one of the best analyzed examples because several significant and granitic plutons that penetrated the former two (e.g., Geological and classic concepts in geology and geophysics were proposed on Survey of Japan, 2003). Such intense deformation, associated regional the basis of direct observations from Japan. These include the Wadati metamorphism, and granitic intrusion confirm that the Japanese (-Benioff) plane (Wadati, 1935), volcanic front (Sugimura, 1965), paired Islands belong to a segment of a long-existing old orogenic belt, as metamorphic belts (Miyashiro, 1961), Pacific-type orogeny (Matsuda traditionally discussed by many geologists in terms of geosyncline or and Uyeda, 1971), structure of modern/ancient accretionary complexes of plate tectonics. In plate tectonic terminology, the active magma- (Kanmera, 1976; Taira et al., 1989; Matsuda and Isozaki, 1991), anatomy tism, high seismicity, and frequent crustal faulting in modern Japan of accretionary orogen (Isozaki and Maruyama, 1991; Isozaki, 1996), prove that the islands currently form an active subduction-related and wedge extrusion of high-pressure metamorphic unit (Maruyama orogenic front. et al., 1996), among other themes. From a historical viewpoint, the geological studies in Japan prior to the 1980s are categorized in the “non-science phase" and “colonial science phase" in the framework of typical pattern in transplanting * Corresponding author. scientific ideas from Western Europe to the rest of the world (Basalla, E-mail address: isozaki@ea.c.u-tokyo.ac.jp (Y. Isozaki). 1 Currently at Department of Earth Science and Astronomy, The University of Tokyo, 1967). However, there were several exceptional pioneering geologists Komaba, Meguro, Tokyo 153-8902,Japan. (e.g., Miyashiro, 1961; Matsuda and Uyeda, 1971; Uyeda and Miyashiro, 1342-937X/$ - see front matter ? 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/jgr.2010.02.015 Y. Isozaki et al. / Gondwana Research 18 (2010) 82-105 83 1974; Kanmera, K., 1976) who proposed independent new ideas of their analysis demonstrated the profound significance of tectonic erosion own. During the 1980s, nearly a decade after the formulation of plate in subduction-related orogen. The latter two new aspects together tectonics, a major conceptual turnover occurred in the tectonic studies imply that the traditional understanding of the uni-directionally e n n d oceanward growth of ACs is not valid, instead emphasize that tectonic coupled with the improved knowledge on blueschists, ophiolites, and erosion is effective in shaping juvenile continental crust as well as granites in the plate tectonic framework (e.g., Maruyama and Seno, accretion. 1986; Taira et al., 1989; Isozaki et al., 1990; Matsuda and Isozaki, 1991). This contribution begins with a brief review of the overall geotectonic In every step of the previous re-evaluation of geotectonic subdivision in framework of the Japanese Islands, together with the main aspects of the Japan, progress in higher-resolution chronology always played the main ca. 600 million year long tectonic evolution. We then explore new role. The subdivision of Japan prior to the 1970s was largely based on perspectives of current studies on geotectonics of the Japanese Islands, ages of megafossils (bivalve, ammonoid, brachiopod, coral etc.) plus and discuss their geological implications. fusulines (e.g., Kobayashi, 1941; Minato et al., 1965). During the 1980s, research on high-resolution microfossil (conodont and radiolaria) 2. Tectonic framework drastically changed the definition of geotectonic elements by identifying ancient accretionary complex (AC) units (e.g., Isozaki and Matsuda, 2.1. Current plate boundaries 1980; Ya0 et al., 1980; Yamato-Omine Research Group, 1981; Ishiga, 1983; Matsuoka, 1992), while the regional K-Ar radiometric dating/ Japan is located at the junctions of four distinct plates; i.e. the mapping enabled a more clear definition of the weakly metamorphosed Eurasian, Pacific, North American, and Philippine Sea plates (Fig. 1). AC units in the 1990s (Isozaki and Itaya, 1991; Isozaki et al., 1992). NE Japan (northern Honshu + Hokkaido） belong to the North Summarizing all the results derived from these innovative America plate, whereas SW Japan (Western Honshu + Kyushu) and practical methodology and relevant new concepts, the history of the Ryukyu Islands to the Eurasian plate. The Pacific plate is subducting at Japanese islands was completely rewritten in the 1990s in terms of a rate of 10 cm/yr beneath NE Japan, with its leading slab down to the plate tectonics (e.g., Isozaki and Maruyama, 1991; Isozaki, 1996, 660 km depth under Beijing, China, as detected seismologically. Japan 1997a, 1997b; Maruyama, 1997). The significant aspects of this s e summary include: 1) the origin of Japan in the South China (Yangtze) Japan, marks the western end of the Pacific plate on surface. The margin, 2) tectonic conversion from a passive margin to an active Philippine Sea plate is subducting from SE to NW with 4 cm/yr under margin, 3) subhorizontal accretionary growth with downward SW Japan along the Nankai trough (trench) and Ryukyu trench off SW younging polarity, 4) episodic ridge subduction and the related Japan. formation of regional metamorphic/granite belts, and 5) microplate NE Japan collides against SW Japan along the over-2000 m-high tectonics (back-arc opening, arc-arc collision, and fore-arc sliver). mountains called Japan Alps that stretch across the arc in central part These aspects in turn offered a general geotectonic context for of the main island (Fig. 1) where the compressional uplifting has describing and understanding subduction-related (previously called continued since 0.5 Ma. There are two TTT (trench-trench-trench)- Cordilleran-type, Pacific-type, or Miyashiro-type) orogenic belts in the rest of the world. Notably, these concepts were not imported from Tokyo to the south, that are constantly shaking the ground of the the advanced countries in the west, but originated from Japan on the nation's capitol. To the north of Ryukyu arc, extensional tectonics has basis of domestic data. In other words, the tectonics study in Japan initiated to form a back-arc basin within East China Sea since 5 Ma stepped out of “colonial science phase" and entered the next stage along the Okinawa trough off the Ryukyu Islands that is propagating termed the “independent science tradition phase" by Basalla (1967). from south to NE towards central Kyushu. The 1990s also witnessed the geotectonic explanation of the Japanese Islands in the context of the repeated assembly and breakup 2.2. Accretionary front of supercontinents since 1.9 Ga (e.g. Isozaki and Maruyama, 1991; Maruyama et al., 1997), in particular with respect to the history of South One of the characteristic geologic phenomena in modern Japan is China continental block that originated from the Proterozoic supercon- the formation of AC on the leading edge of Eurasian plate. The seismic tinent Rodinia. In the mean time, several ambiguous interpretations reflection profiles in SW Japan have clarified that trench-fill sediments were applied to Japan (Ichikawa et al., 1990); e.g., the “suspect (exotic) have been under strong compressional deformation, i.e. folding and terrane" concept (e.g. Jones et al., 1977; Coney et al., 1980) coupled with thrusting with an oceanward vergence, and these sediments have large strike-slip dislocations for over thousands of kilometers (e.g., Taira been successively incorporated into ACs with downward younging et al., 1983; Yamakita and Otoh, 2000; Sakashima et al., 2003). Such polarity (Kuramoto et al., 2000). The main thrust corresponds to a anachronistic return of the “colonial science", however, could not decollement zone between the accretionary wedge and the subduct- explain the essentially subhorizontal nature of the orogen in Japan, and ing Philippine Sea plate. Although the majority of accretionary wedge those models were eventually proved to be unrealistic. is composed of trench-fill turbidites, a minor amount of oceanic rocks The fundamental aspects on the orogenic evolution of Japan are often accreted together with the former. mentioned above have been widely accepted for nearly two decades, Japan has basically grown oceanward in map view by the and the geotectonic studies became apparently dormant for a while. formation of ACs since 500 Ma (Late Cambrian), but its growth has not been continuous, instead it was intermittent (Isozaki and introduced from various corners of geology/geophysics. These include Maruyama, 1991). In contrast to SW Japan that is enriched in the 1) detailed seismic images of the deep crust and mantle beneath Japan Cenozoic ACs along Nankai trough, NE Japan is characterized by a (Hasegawa et al., 2009; Zhao, 2007; Sato et al., 2005), 2) protolith considerably small amount of post-Miocene AC next to Japan Trench chronology of regional metamorphic rocks using detrital zircon (Aoki owing to the active tectonic erosion (von Huene and Scholle, 1991). et al., 2007, 2008), and 3) provenance analysis by detrital zircon n ir (s " n 2.3. Arc volcanism izontal structure of the deep crust of Japan and heterogeneity of sub- arc mantle. The results from the second approach warranted a re- Volcanic activity is another characteristic feature of modern Japan, definition of the regional metamorphic belts in Japan. This is again and so was in the past. The chain of currently active volcanoes, in related to high-resolution dating of orogenic elements and demand particular those forming the volcanic front on the Pacific side, runs re-definition of regional metamorphic belts in Japan. The third fundamentally parallel to the active trench. Those on the volcanic 84 Y.Isozakiet al./GondwanaResearch 18(2010)82-105 Eurasia plate platey NorthChina Pacificplate D 10cm/yr >4cm/yr 2.9-1.9Ga Philippine Yangtze Seaplate B Shibao 1.8Ga Q S.China 600Maophiolite (Cathaysia) TTTK ① 14Ga PHS TriassicUHP-HP 600Maophiolite KoreanPen. collisionzone Japan trench N.China N.China S.China A TriassicUHP-HP Cathaysia. B collisionzone Yangtze PhilippineSea plate 1.0-0.9Ga Shibao orogen Phanerozoic accretionary complex Fig. 1. Geotectonic setting around the Japanese Islands in East Asia. The island-arc chains around Japan comprise the Kuril,Japan, Ryukyu and Izu-Bonin-Mariana arcs formed by the subducted by the Pacific plate, whereas SW Japan (+ Ryukyu arc) belongs to the Eurasia plate that has been subducted by the Philippine Sea plate. The Japan Sea behind NE and SW Japan represents a Miocene rifted back-arc basin. The Okinawa trough behind the Ryukyu arc is a nascent rift zone to form another back-arc basin in future. The main part of the Japan and Ryukyu arcs has a strong geologic ink to South China, because Japan has developed along an active continental margin of South China since the Early Paleozoic by the long-term subduction processes from the Pacific side.South China block is composed of two parts,the Yantze block on the northwest and the Cathaysia block on the southeast, that became united by a ca. 1.0 Ga collisional suture called the Shibao orogen (Li, 1999). The ca. 600 Ma ophiolite unit in Japan probably represents a remnant of pre-existing passive margin complex prior to the initial subduction regime.South China merged with North China by the 240-230 Ma collisional event that produced the Dabieshan-Sulu suture characterized by an ultrahigh-pressure (UHP) metamorphic belt. Note the subhorizontal superposition of North China above South China (see the profile) along the collision suture with the UHP unit. The eastern extension of the Dabieshan-Sulu suture can be traced to the Imjingan belt in northern Korean peninsula, and its smaller fragments further to the east are alsorecognized sporadically in SW Japan mostly along the Japan Sea margin. Judging from the Shibao orogen in central Korean peninsula, proto-Japan (at least during the Paleozoic) had a direct link to the Cathaysian part of South China. front occur approximately above 100-120 km above the relevant side) and ACs (on the Pacific side). This regional arrangement subducted slab, regardless of the subduction angle, suggesting a suggests that the Japanese Islands form a segment of the Mesozoic pressure-dependence for the calc-alkaline magma generation in deep subduction-related orogen that grew along East Asian margin at least subduction zone. High-resolution P-wave tomographic images re- since the Jurassic time. The formation of eastern half of Asia started cently documented a clear image of the subducting cold Pacific slab roughly since the Triassic by the initial collision/amalgamation and a rising flow of mantle in the overlying mantle wedge with processes between Indochina, South China, and North China (Sino- volcanoes right above (Zha0 et al., 2009; Hasegawa et al., 2009, 2010). Korean) blocks, together with other minor blocks (Maruyama et al., The low-velocity perturbation in the hanging wall corresponds to the 1989). As to the earlier evolutionary history of Japan during the secondary mantle convection rising up to intersect the arc crust where Paleozoic, South China likely played a relatively important role for the volcanic front develops (Maruyama et al., 2009). Minor volcanoes the orogenic growth rather than North China (Fig. 1), as indicated by appear behind the front, called the second chain of volcano on the the physiological continuity to Kyushu and by the similarity in the Japan Sea side. Paleozoic faunal composition (e.g., Nakazawa, 1991). North China joined later after its collision with South China in the Triassic (230- 2.4. Japan in East Asia 240 Ma), an event that generated the ultrahigh-pressure (UHP) metamorphic belt of the Dabieshan-Sulu suture. The eastern exten- In East Asia, all the way from Kamchatka (Far East Russia) to West sion of this suture is traced up to the Korean peninsula (i.e., the Philippines, a parallel alignment constantly develops between the Imjingan belt; Cluzel et al., 1990; Oh, 2006). However, further exten- Jurassic-Cretaceous paired belts of granite batholith (on the continent sion to the east has been unclear, although the 230 Ma medium- Y. Isozaki et al. / Gondwana Research 18 (2010) 82105 S8 pressure metamorphic rocks in central Kyushu is a possible candidate Cenozoic orogenic elements, i.e., highly-deformed sedimentary rocks for this extension. It is noteworthy that South China is structurally and regional metamorphic rocks, that align roughly parallel to the positioned below North China, and that the collision-related UHP general trend of the Japanese Islands (Fig. 2). Nearly 80% of these metamorphic belt between them is in subhorizontal attitude (Fig. 1 basement rocks in Japan consist of ancient ACs, whereas some profile). The southeastern corner of the Korean peninsula represents a fragmentary non-accretionary elements sporadically occur mainly large-scale klippe of North China that sits on the UHP-bearing suture along the Japan Sea sidle. Most of these AC elements underlie NE and and South China. SW Japan, and also in part Ryukyu Islands. On the other hand, the Izu The juvenile Japan during the Late Proterozoic to Early Paleozoic Bonin arc and Eastern Hokkaido (a part of Kurile arc) represent likely developed in the vicinity of South China.South China consists of two parts; i.e., Yangtze block on the northwest and the Cathaysia formed above the ca. 10o Ma oceanic lithospheres by the subduction block on the southeast, that were merged by the ca. 1.0 Ga collision of another oceanic plate. Through the Miocene collision against the that left the Shibao orogen between them (Li, 1999). The ca. 600 Ma major arc systems along East Asia, these two exotic systems became ophiolite unit in Japan likely represents a remnant of pre-existing incorporated into Japan islands. Owing to their exotic nature, the passive margin complex prior to the initial subduction regime. As following description excludes Izu-Bonin and E. Hokkaido. South China with the Shibao orogen is exposed in the central Korean Two distinct groups of geologic units are discriminated in terms of peninsula (Fig. 1), the Japanese Islands likely had more intimate plate tectonics-based orogenic concept: 1) units formed along the Early connection with the Cathaysian part of South China rather than the Paleozoic passive continental margin mostly of South China block and Yangtze part. later detached from the mainland, and 2) units formed in-situ along the Pacific subduction-related active margins (Table 1; Isozaki and 3.Major geotectonicunits and boundaries Maruyama, 1991; Isozaki, 1996). A very minor group of units derived from the continent-continent collision zone between South China and By removing the Neogene-Quaternary sediments/volcanics and North China blocks is also associated as a possible eastern extension of post-jurassic granitic intrusions, we can perceive that the surface the DabieshanSulu suture. These smaller pieces have a significant crust of Japan is mainly composed of the Late Paleozoic to Early NorthAmericanplate Rebun Japanbasin 6 301 Eurasiaplate 300 Tk(Sh+Sb) Yamatotai 302 794 Kitaoki IPsubduction complex Yamatobasin mJ-eK subduction complex Okino Fig.5 mplex 297 collapsed seamount Pacificplate Kyushu-PalauRidge Low-P/T type metamorphic belt (arc) 442 High-P/T type metamorphic belt Continental block(N.China,S.China) Fig. 2. Geotectonic subdivision of the Japanese Islands (modified from Isozaki and Maruyama, 1991). The geological units are abbreviated as follows: SW Japan: Rn: Renge belt; Sn: Sangun b.: Ak: Akiyoshi b.; Mz: Maizuru b.; UT: UItra-Tanba b.: M-T: Mino-Tanba b.: Ry: Ryoke b:; Sb: Sanbagawa b.: Ch: Chichibu b.: Sh: Shimanto b, NW Japan: Ht-Tk: Hitachi- Takanuki b:; Gs: Gosaisho b; MM: Matsugadaira-Motai b.; SK; Southern Kitakami b.: NK: Northern Kitakami-Oshima b.; Kk: Kamuikotan b.: Hdk: Hidaka b.; T: Tokoro b; Nm: Nemuro b. The accretion is active along the Nankai trough, whereas the erosion of arc material dominates along the Japan trench. Their mutual difference likely comes from the contrasting useq sue-xoeq au wuoy uede[ Ms jo alyod squopajo8 au og s B! aag ajeld ogoed piu pio/pro snsian aned eas audd agonp ou/uno. ar atepd Suonpqns jo aunjeu Japan Sea) margin to the active trench (Nankai Trough). 86 Y. Isozaki et al. / Gondwana Research 18 (2010） 82-105 Table 1 et al., 1990; Matsuda and Isozaki, 1991). Although all ACs have more Major orogenic elements and criteria of the active and passive margins. or less the same rock association, they can be discriminated strictly 1) Passive margin elements from each other solely in terms of OPs, thus mutual boundaries Older granite-gneiss complex between neighboring ACs are clearly defined. Detailed field mapping Fore-arc ophiolite of ACs coupled with OPS analysis clarified that all AC units in Japan Rift-ill and continenta rise sediments Occur as thin subhorizontal fault-bounded geologic bodies, i.e., nappe, Remnant of LIP 2) Active margin elements and that on the whole they show a clear downward (and oceanward) Arc plutonics/volcanics younging polarity (Isozaki and Itaya, 1991; Isozaki and Maruyama, Accretionary complex (AC) 1991). Fore-arc sediments of the Ordovician to Cenozoic cover these High-P/T metamorphosed AC ACs, high-P/T meta-ACs, and granitic intrusions. Fore-arc sediments Criteria 4. Brief history of the Japanese Islands For sedimentary rocks (Mega, micro-)fossils The essential story of the ca. 7o0 million year-long geotectonic Lithofacies evolution of Japan is briefly summarized here. The history is divided Stratigraphy (normal, OPS) For metamorphic rocks into 3 distinct stages: 1) passive continental margin stage during 700- Mineral assemblage 500 Ma, 2) active continental arc stage during 500-20 Ma, and Radiometric age 3) active island arc stage since 20 Ma (Isozaki, 1996, 2000; Maruyama, Protolith type/age 1997; Fig. 3). For igneous rocks The birth of Japan was related to the ca. 7o0 Ma breakup of the Late Rock series Geochemistry (trace, REE etc.) Proterozoic supercontinent Rodinia, in particular, its fragment called Radiometric age South China block. The breakup of Rodinia is mirrored in the birth of Pacific (Panthalassa) Ocean. Proto-Japan formed a part of the passive continental margin of South China for the first 200 million years until Passive continental margins are generally characterized by pre-rift mid-Cambrian ca. 500 Ma. Around 500 Ma, a tectonic conversion older granite-gneiss complexes, fore-arc ophiolite, fragment of large occurred from a passive to an active margin by the initiation of igneous province (LIP), rift-filling sediments, and extremely thick subduction from the Pacific side. The subduction-controlled tectonic continental rise sediments. The Oki belt and the Hida belt in SW Japan regime continued for 500 m.y., although a small-scale rifting of are composed of granite-gneiss complexes that were penetrated by continental margin occurred in the Miocene to open a back-arc basin the Jurassic granites (Fig. 2). The Jurassic batholiths in East Asia occur (Japan Sea) that modified Japan from a continental arc to an island arc as commonly along the present Asian margin from South China to we see today. Kamchatka but not in Japan except the Oki and Hide belts (Takahashi, When we extrapolate the current plate movement to future, all the 1983). The Jurassic granites in the Oki and Hida belts probably formed dispersed continents except Antarctica will merge together once segments of subduction-related batholith beneath the coeval volcanic again to form a supercontinent of the next generation, Amasia arc along the pre-Japan Sea continental margin of East Asia, (Maruyama et al., 2007). Given the present plate motions working representing fragments of older inboard elements with respect to exactly the same in future, the following processes are expected to the accretionary units described below. The oldest ophiolite in Japan occur around Japan. In the coming 50 million years, Australia will occurs as a small block at Nomo in western Kyushu, and was dated ca. collide against the eastern margin of Eurasia by crushing many 583 Ma (Ilgi et al., 1979). As all subduction-related units are younger archipelagos of Papua New Guinea, the Philippines, and Japan. than 500 Ma, this likely represents a fragment of the oldest oceanic Furthermore, North America will collide against Eurasia in 250 million crust formed by the breakup of South China from the supercontinent years by closing the Pacific Ocean and forming a new supercontinent. Rodinia prior to the appearance of the long-lasting subduction system. Thus, in terms of the Wilson Cycle of a major ocean, the geotectonic No fragments of LIPs related to the initial breakup of Rodinia nor rift- history of Japan, for 700 plus 250 million years from its birth to its related, continental-rise sediments have been identified yet in Japan. demise, represents a mirror image of the entire history of the Pacific Elements of active continental margin are dominated mainly by Ocean (Maruyama, 1997). ACs (including high-P/T metamorphosed parts） and calc-alkaline granitic plutons penetrated into the former (Fig. 2). In Japan, the 5. New aspects during the last decade oldest granitic unit is ca. 500 Ma, whereas the oldest AC (high-P/T metamorphosed) is ca. 450 Ma. The youngest granite and AC exposed 5.1. Deep seismic profiles on land in Japan are both of Miocene age. Granitic batholiths show a zonal pattern in Japan; however, due to their complicated mosaic 5.1.1. Upper mantle Recent seismic studies using the P-wave tomography identified sharp images of the old Pacific slabs under Japan (e.g., Fukao, 1992; distribution pattern of ACs, in particular in SW Japan. Zhao et al., 2007; Hasegawa et al., 2009). The extension of the Pacific On the basis of formation ages, the ACs were subdivided into Plate is traced continuously down to the mantle boundary layer (MBL; the following 9 units (traditional zone names in bracket): 420 Ma 410-660 km depth range) (Fig. 4A). Such a deep penetration of the high-P/T meta-AC (Oeyama belt), ca. 340 Ma high-P/T meta-AC slab proves that the subduction of the Pacific plate must have (Renge belt), 260 Ma AC (Akiyoshi belt), 240 Ma high-P/T meta-AC e oe qqod e ooi s ns ne uonp s e o pn (Suo belt), 220 Ma AC (Ultra-Tanba belt), 160Ma AC (Mino-Tanba old trench located at more or less the same position as the present belt including the outlier called the Chichibu belt), 120 Ma high-P/T one. The vertical profile between northern China and the Japanese meta-AC (Sanbagawa belt), 80Ma AC (Northern Shimanto belt Islands particularly indicates that a mass of flat-lying (2000 km wide) including 60 Ma high-P/T meta-AC), and Paleogene-Miocene ACs stagnant slab exists over the MBL in East Asia (Zhao et al., 2007). The (Southern Shimanto belt). Accreted material in an AC is the key to high-V anomaly and low-V anomaly beneath the Pacific slab suggests estimate the precise age of subducted slabs, as ocean plate that the catastrophic break-up of stagnant slabs may have occurred to stratigraphy (OPs) records the travel log of the relevant ocean plate cause the mantle overturn from the lower mantle (Maruyama et al., from its birth at a mid-oceanic ridge to its demise at a trench (Isozaki 2007b). Y. Isozakiet al./Gondwana Research 18(2010)82-105 87 Phanerozoic Neoproterozoic Age Paleozoic Mesozoic Cenozoic (100 Ma) 6 5 4 3 2 0 -1 1. 12. 3. 4. 5. 6. Birth Growthstart Collision Arc-isolation Growth-end Merge rifting-Pacificopentectohicinversion (Yangtze/Sino-Korea) JapanSea Australia-collisionN.A-colision major events inJapan 1.Atlantic-type Il.Pacifi-type active contin.margin Ill.Coalescence passivemargin XxX -Oeyama Hida AC riftzone+ Mino-Tanba Collision-type HP-AC continental Kur orogenicproducts rise deposits w/UHProcks E 190-150 First subducted subduction age of subducted oceanic start 100 platesand theirage 150 at trench (m.y.) R ridge subduction at trench Nomo*/ Renge* Farallon/ /Kula/PacificPhil.Sea Rodinia Gondwana /lapetus Pangea expand Tethys/ expand superocean Mirovia proto-Pacific Panthalassa shrinkPacific superplume ★Pacific ★Africa ★ Fig.3. Chronicle of geotectonic evolution of Japan from its birth at ca. 700 Ma to its demise at ca. 200 million years after the present (modified from Isozaki, 2000). Geologic proceses in Japan, such as formation of AC unit, exhumation of high-P/T meta-AC unit, emplacement of granitic batholith etc. are summarized with respect to the history of subducted oceanic plates and supercontinent/superocean history.Age of the subducted oceanic plate at trench was estimated from ocean plate stratigraphy (OPs) analysis of AC units. Note that the geotectonic history of Japan can be divided roughly into 3 major stages: 1) the passive margin stage (700-520 Ma), 2) the active continental margin stage (520 Ma to -50 Ma) with the island arc sub-stage,and 3）the collision and coalescence stage(-50 Ma to-200 Ma).Byextrapolating the current regime of plate motions,what is expected next is the stage of final colision (ca.50-2o0Ma after the present of the Australian and North American continents against Eurasia to form Amasia,a future supercontinent. This geotectonic chronicle of the Japanese Islands corresponds essentially to the Wilson cycle of the Pacific Ocean that appeared around 700 Ma and will disappear in another 200 million years, as marked by superplume activities in mantle.The continental growth in Japan segment in East Asia has occurred solely during the second stage when subduction-related magmatism produced scale erosion and mass wasting (= material recycling) from the pre-existing arc crust. Net addition of new crust was culminated several times when mid-oceanic ridges subducted beneath Japan margin; however, tectonic erosion often took away the pre-existing fore-arc crust. The seismic images beneath NE Japan demonstrate a clear SW Japan is clearly divided into two contrasting domains separated by arrangement of earthquake epicenters within the Pacific slab and in the Median Tectonic Line - MTL (Sato et al., 2005; Ito et al., 2009): a the arc crust of NE Japan (Nakajima et al., 2001; Fig. 4B). It is granite-batholith-dominant part on the continent side and the AC (+ noteworthy that earthquakes occur not along the top surface of the meta-AC)-dominant part on the ocean side. The former is composed of a subducting Pacific slab but along the “double seismic zones" within 15-km-thick lower crust of unknown composition, a 20 km-thick the slab; i.e. in the topmost-slab and in mid-slab mantle. These granitic upper crust, and thin roof-pendants of the Paleozoic to Jurassic earthquakes are likely caused by the dehydration embitterment of AC (+ meta-AC) units on the surface. Detailed subhorizontal structure dense hydrous magnesium silicates (Omori et al., 2002, 2009). Water- of these pre-Cretaceous AC and meta-AC units will be described later. rich fluids liberated by the dehydration reaction may lead to local and On the ocean side, the crust is composed solely of the Late Mesozoic to Cenozoic AC and meta-AC units that are characterized by this likely causes earthquakes within the slabs. On the other hand, subhorizontal stack of multiply fault-bounded units (previously called deep earthquakes originated in MORB crust of subducted slabs are nappes) with a very gentle dip angle towards the Eurasian continent dominated down to 50-60 km depth, but limited to around 100 km (Fig. 5). Particularly impressive is the thick pile of Late Cretaceous to depth, by the drying hydrated MORB crust through dehydration Tertiary ACs and meta-ACs of the Shimanto belt together with much reactions. These essentially reflect the circulation pattern of fluids younger sub-seafloor ACs that occupy most of the fore-arc crust. beneath fore-arc, and may constrain the domain of subduction- Furthermore, recent geophysical analyses (Sato et al., 2005; Ito et al, related high-P/T metamorphism. into the continent side even beyond the surface transect of the MTL, as 5.1.2. Crust previously predicted according to the surface geology (Isozaki, 1988; As to the arc crust, earthquakes occur dominantly only in its upper Isozaki and Itaya, 1991; Isozaki and Maruyama, 1991). In short, the half (Fig. 4B). This is likely due to the depth-dependent ( = temperature- essentially subhorizontal internal structure of the Sw Japan crust was dependent) brittle-ductile transformation of coarse-crystalline granitic geophysically confirmed in accordance with its surface geology, and rocks. The boundary between the seismic and aseismic domains this clearly disproved the once popular but unrealistic strike-slip o n n sn fault-controlled tectonics in SW Japan (e.g., Taira et al., 1983; which granitic rocks behave in a brittle manner (Shimamoto, 1986), Yamakita and Otoh, 2000; Yao, 2000). hence they are capable of triggering earthquakes. In other words, this At the Nankai Trough, in particular, high resolution seismic reflection study clarified the detailed image of on-going frontal crustal detachment". accretion processes at a modern active trench (Kuramoto et al., 2000; The across-arc crustal profile of Japan, in particular in SW Japan, was Fig. 5). Combined with on-land geological observations, the pattern of for the first time visualized by seismic analyses using seismic reflection, overall accretionary growth was documented for the leading edge of vibroseis, and explosive-dependent experiments (Fig. 5). The crust of the Asian continental plate. 88 Y.Isozaki et al./Gondwana Research18(2010) 82-105 A Changbai Datong trench trench A'B B A 01 500 1000 1500 2000 2500 1% 0% +1% B D-D' E-E' G-G' H-H Fig. 4. Seismic tomography beneath East Asia with across-arc profile of NE Japan. A: Vertical cross sections of whole mantle P-wave tomography of East Asia in latitudinal sections across theJapan arc Zhaoand Ohtani 2009).Red and blue color denotes slow and fastvelocities,respectively.White dots and twosolid lines indicate arthquakes within10kmf each profile, and discontinuities of 410 km-deep and 660 km-deep with mineral phase transition. Note the flat-lying high-velocity anomaly in the mantle beneath East Asia that to older subducted slabs sinking from the mantle boundary layer to the 29oo km-deep core-mantle boundary.B: Across-arc vertical section of S-wave velocity perturbations in NE Japan (Nakajima et al., 2001).Solid black bar and open circles represent land area and earthquakes, respectively. Orange-yellow colored low-V zone develops in the upper part of the mantle wedge,probably representing a counter-flow of mantle convectiontriggered by the Pacific slab subduction Rising counter-flow of mantle intersects with arc crust to release arc magma beneath volcanic arc, by melting mantle rocks due to suffcient water form Pacific slab and to pressure decrease. Note the double-seismic zone developed within the subductingslab. 5.2.Detritalzirconchronology1:protolithageofmetamorphicrocks In SW Japan, at least four distinct high-P/T metamorphic belts were traditionally discriminated; i.e. 450 Ma Oeyama, 340 Ma Renge, 240 Ma Geotectonic definition of high-P/T metamorphic units has been Suo, and 120 Ma Sanbagawa units. The main age constraints were fossils mainly based on metamorphic facies series and ages of protolith/peak in protoliths and radiometric data from metamorphic minerals. As the metamorphism/cooling. All the subduction-related metamorphic belts, best-documented example, the world-famous Sanbagawa metamor- however, are similar to each other in metamorphism per se, thus their phic rocks was regarded as the metamorphosed earliest Cretaceous (ca. chronology is the most critical factor in distinguishing metamorphic 140 Ma) AC that was subjected to peak metamorphism around 120 Ma units. The chronology of metamorphic units has been conventionally (Early Cretaceous) and exhumated during 110-80 Ma (e.g., Itaya and constrained by three distinct criteria: 1) pre-metamorphic protolith age Takasugi, 1989; Is0zaki and Itaya, 1990; Takasu and Dallmeyer, 1990; by fossils, 2) age of metamorphism by radiometric dating, and 3) age of Okamoto et al., 2004). terminal surface exposure/erosion by geological evidence (e.g., uncon- Recent progress in zircon chronology using ICP-MS analysis, however, formity and conglomerate) as demonstrated for the Sanbagawa and demands a major re-consideration of the protolith age of metamorphic Franciscan high-P/T units (e.g., Isozaki and Itaya, 1990; Okamoto et al., units. Detrital zircon grains contained abundantly in psammitic schists 2004; Sedlock and Isozaki, 1990; Isozaki and Blake, 1994). (meta-sandstone) retain primary isotopic signals for U-Pb dating Y. Isozaki et al./Gondwana Research 18(2010) 82-105 89 Largethrust First out-of slicezon trustzor I-KTL -MTL Maizuru-UItraTanbaB. Landwarddipping Imbricate Suo B,/Ming-Tanba B.! AKL! reflectorszone thrustzon Seto inlandSea San-inB. +Sanyo B. RyokeB. Nankai Trough 10 10 20 Lowercrust 30 sedimentaryrocks 30 Sanbagawa meta-AC Chichibu composite belt seismogeniczone 40 oceaniccrust 50- Mantle MioceneAC 50 Shimantometa-AC m Wadati-Benioffplan Philippine Sea Plate 40km T Fig.5. Geotectonic profile of SW Japan from the Japan Sea coast to the Nankai Trough that depicts the crustal structure between the volcanic front and trench of the matured island separated by the Median Tectonic Line (MTL). The former is composed mostly of arc-related granitic batholiths penetrating into the pre-Cretaceous AC and meta-AC units that remain as roof pendant. The latter is composed mostly of subhorizontal stack of the post-Jurassic AC and meta-AC units with clear downward younging polarity. The pre-Cretaceous AC and meta-AC units on the Japan Sea side also possess the same structure,suggesting an overall growth pattern of AC and meta-AC units in an oceanic subduction-related orogen. The juxtaposition of the two contrasting domains, i.e., the granite-dominant domain and the non-granitic AC dominant one, were made by the activation of the low-angle MTL that likely formed as a mid-crustal detachment in the fore-arc during the Miocene when the back-arc (Japan Sea) spreading occurred. regardless of metamorphic overprints. On the basis of a large number of in ancient arc-trench system. This method is more effective than the detrital zircon ages obtained by ICP-MS, Aoki et al. (2008) pointed out conventional clast-analysis of conglomerates because relatively hard that the traditional “Sanbagawa belt" in Shikoku needs subdivision into (resistant to weathering) zircon grains are abundant in sandstone and two independent units: 1) the Sanbagawa unit sensu stricto formed as AC at ca. 140-130 Ma and high-P/T metamorphosed at 120-110 Ma, and 2) a by-grain with the advanced micro-scale chronological techniques. newly recognized unit formed as AC after ca. 80 Ma and high-P/T By using ICP-MS analysis, Nakama et al. (in press-a) recently metamorphosed at ca. 60 Ma (Fig. 6A). As the 120-110 Ma schists and the 60 Ma schists have more or less the same sets of metamorphic Paleozoic to Cenozoic sandstones in Japan, showing the secular change mineral assemblages of the so-called “high-pressure intermediate group", of zircon population in terrigenous clastics deposited around Japan it has been difficult to discriminate them solely with conventional petro- (Fig. 7). The pattern clearly shows that the source areas of terrigenous logical dataset. However, the latest high-resolution studies confirmed clastics have changed drastically during the over-500 m.y. history of that these two are not continuous with respect to the distribution area, the Japanese Islands. Except for minor contribution from recycled protolith age, and the timings of peak metamorphism and exhumation. grains, the dated zircon grains are clustered into 7 distinct groups; i.e., In terms of AC-forming age, the 60 Ma blueschist unit corresponds to the 2500-1000 Ma (Paleo- to Mesoproterozoic), 1000-800 Ma (Neopro- non-metamorphosed AC unit of the traditional N. Shimanto belt terozoic), 520-400 Ma (Cambrian-Silurian), 280-210 Ma (Permian- (Tsutsumi et al., 2009), and at the same time, it represents the youngest Triassic), 190-160 Ma (Jurassic), 110-90 Ma (mid-Cretaceous), and blueschist unit in SW Japan (Fig. 3). These support the results from grain- 80-60 Ma (Late Cretaceous-Paleogene), as illustrated in Fig. 7. These 7 composition analysis of meta-sandstones of the "Sanbagawa belt" groups essentially represent ancient provenances of terrigenous (Kiminami et al.,1999). clastics that were transported to and deposited in Japan. In addition, recent analyses of micro-inclusions in metamorphic The major points to be noted in this connection are as follows: minerals by nano-SIMS clarified that the conventional metamorphic 1) terrigenous flux from 2 major neighboring continents (North and zones reflected not a progressive metamorphic facies series toward South China blocks) was very small, far minor than previously expected, the peak metamorphism but a retrogressive one associated with 2) the mid-Paleozoic to Early Triassic sandstones contain numerous hydration processes during the exhumation. Aoki et al. (2009) has zircon grains of ca. 520 to 400 Ma age, 3) the Triassic to Jurassic demonstrated multiple metamorphic overprints during the post-peak sandstones are dominated by ca. 290-210 Ma grains, 4) the Proterozoic metamorphism exhumation stage onto the primary eclogitic signa- grains occur almost solely from the Middle-Late Cretaceous sandstones, ture, and confirmed again that two distinct blueschist belts of and 5) post-mid-Cretaceous sandstones lack older grains but contain different exhumation ages exist side by side within the traditional abundant Late Cretaceous-Paleogene zircon grains. In short, the "Sanbagawa belt" (Fig. 6). For further details of the revised definition sandstones deposited during the 600 m.y. history of Japan can be of high-P/TAC belts, such as the well-known “Sanbagawa belt", refer a d jo sn u s s e ni p to Aoki et al. (2007, 2008). 1) Paleozoic to Triassic sandstone with 520 to 400 Ma clastics, 2) Jurassic to Late Cretaceous sandstone with 290-150 Ma clastics, and 3) post- 5.3. Detrital zircon chronology 2: provenance of clastics mid-Cretaceous one with 100 Ma clastics (Fig. 7). The data indicate that the major provenance regime around Japan changed twice; first at ca. Detrital zircon chronology has also been employed as a potential 200 Ma (Triassic-Jurassic boundary) and second at around 90 Ma (Late Cretaceous). It is noteworthy that the Paleozoic and Triassic sandstones in Japan hardly contain terrigenous clastics from the two neighboring The mineral zircon forms commonly in felsic igneous rocks, in major continental blocks, i.e., North China underlain extensively by the particular plutonic rocks such as tonalite/tronjemite/granite suite, Archean to Early Proterozoic rocks and South China with the Late that often form large batholith belts along active continental margins. Proterozoic basement, despite their proximity to Japan. On the other In arc-trench settings, these subduction-related batholith belts com- hand, the acute sedimentary influx of Proterozoic grains was recorded monly become potential provenances that supply abundant terrigenous solely in the Middle-Upper Cretaceous sandstones. This suggests that a clastics (including zircons) to active trench and its environs. Thus by short-term episodic event occurred in the major provenances and the checking the age spectrum of detrital zircon grains in sandstone from distributary systems of terrigenous clastics in Cretaceous East Asia. subduction-related orogenic belts, it is possible to monitor the Nakama et al. (in press-b) also clarified two significant points. exposure/erosion status of continental crust and sedimentary regime 1) The “so-called Ordovician unit" of the Hida marginal belt contains 90 Y.Isozakiet al./GondwanaResearch 18(2010)82-105 volcanic tuff with igneous zircon grains of ca. 472 Ma (Early have been the core of the geotectonic subdivision of Japan since the Ordovician). This marks the oldest non-metamorphosed sedimentary 1990s (Isozaki and Maruyama, 1991; Isozaki, 1996). Nonetheless, unit in Japan hitherto known . 2) The Miocene rift-related sediments some ambiguity remained because some AC units are not fossiliferous, along the Japan Sea coast recorded the termination of provenance link in particular, datable fossils are extremely poor in trench-fill between the mainland Asia and the modern Japan arc with respect to turbidites, i.e. coarse-grained terrigenous clastics (sandstone) that the opening of the back-arc basin, the Japan Sea. For further details of 'u o no d d o the provenance analysis by detrital zircon, refer to Nakama et al. rocks are less fossiliferous, and therefore impossible to constrain the (2010a, b). protolith age, that is, the precise formation age of AC prior to the deep subduction/metamorphism in most of the high-P/T metamorphosed 6. Discussion AC (meta-AC) units. In contrast, detrital zircons in meta-sandstone, in particular in the The new data and models synthesized in this work were acquired core of single zircon crystals, often survive later metamorphic mainly from several key areas in SW Japan and future studies will extend overprint of isotopic composition. As discussed previously, the recent their application to the rest of Japan. Based on the available data, we progress in chronology of detrital zircons from the “Sanbagawa high- attempt below a preview of the possible revised interpretations. P/T unit" has helped in clarifying significant aspects including the division of the world-famous “"Sanbagawa high-P/T belt" in Shikoku 6.1. Redefinition of orogenic units and boundaries into two distinct units (Fig. 6). This warrants a major re-evaluation of the geotectonic subdivision not only in Shikoku (the classic area for The formation age of AC is determined in general on the basis of the geotectonic studies) but also in the rest of Japan. According to the ocean plate stratigraphy (OPS; Isozaki et al., 1990; Matsuda and preliminary survey in the Kanto Mountains and in central Ki Isozaki, 1991), utilizing microfossil ages. The same approach is peninsula by S. Yanai, K., Aoki and their colleagues, this two-fold fundamentally applied also for metamorphosed ACs. Identifying OPS subdivision can be applied to the entire "Sanbagawa belt" that for each AC unit and distinguishing it from the neighboring AC units extends for ca. 10o0 km in E-W direction along the SW Japan arc. In A Shimanto meta-AC MTL Sanbagawa meta-AC BTL .BTL BTL MTL- ChichibuAC BTL ATL BTL Shimanto AC Sanbagawameta-AC(90Ma) Chichibu AC (Jur-E. Cretac.) highest gradepart w/Sanbosan-MikabuUnit ATL:Aki Tectonic Line 2.80 BTL:Butsuzo Tectonic Line ATL Northern Shimanto AC(L.Cretac.) MTL:Median Tectonic Line 5.21 North unit (w/ cover) - East Shi meta-AC(60 Ma) Midle Unit -+-. 7.89 888 7.76 Lower Unit 0 100km 8.00 trench system, i.e., high-P/Tbelt near trench and granite batholith beneath volcanic arc.A: New subdivision of Shikoku (compiled by S. Yanai, K.Aoki and their colleagues). MTL: Median Tectonic Line,BTL: Butsuzo Tectonic Line,ATL: Aki Tectonic Line.Detrital zircon chronology and age relationship between the protolith deposition and the peak high-P/Tmetamorphism 7.16 2.98 7.28 2.44 the non-metamorphosed AC units of the Chichibu belt and Northern Shimanto belt.B: Two distinct pairs of the Cretaceous-Paleogene arc-trench system in SW Japan (compiled from 3.33 belt. The meta-AC unit was formed and exhumed along the Wadati-Benioff plane beneath the fore-arc next to trench, while the granite batholith formed beneath the volcanic arc.Their formation was likely induced by the ridge subduction. SW Japan suffered ridge-subduction twice during the Cretaceous-Paleogene time (Fig. 3). Y.Isozakiet al./GondwanaResearch18(2010)82-105 Sanbagawahigh-P/Tmeta-AC +Sanyo batholith belt (Low.Cret.) AC (Sanbosan AC + Sanbagawa meta-A uesog (b) Early Cretaceous pair Volcanic rocks 100 km Ryoke belt Sanyo belt Fig.6(con Shimantohigh-P/T meta-AC N. Shimanto (Up. Cret.) AC Up. Cret. fore-arc basin San-in batholith belt belt San-in (Shimanto AC + meta-AC) B 92 Y.Isozakiet al./Gondwana Research 18(2010)82-105 the Kanto Mountains, for example, more than a half of the traditional suggested by the surface field mapping (Isozaki and Maruyama, 1991) "Sanbagawa belt" likely corresponds to the Shimanto meta-AC (e.g., and recent seismic profile (Ito et al., 2009). On the other hand, the Tsutsumi et al., 2009). Shimanto meta-AC is exposed strictly in a tectonic window beneath The recognition of a new geotectonic unit also demands re- the Sanbagawa meta-AC nappe, in a domal domain around the definition of the belt-defining boundary faults not only between the two units of the “Sanbagawa belt", but also among the neighboring mode of occurrence of the two meta-AC units is concordant with the units. In a map view, the northern margin of the classic “Sanbagawa overall structural framework of all Phanerozoic AC units in Japan belt" is clearly demarcated by the high-angle MTL (neo-MTL) that has (Fig. 5) as predicted from the OPS comparison between the activated during the Quaternary in a right-lateral strike-slip sense, Sanbagawa meta-AC and Shimanto AC (Isozaki, 1988). whereas the southern margin is defined as a sharp metamorphic gap with the non- to weakly metamorphosed Jurassic AC unit (of the of the Shimanto belt (more specifically Northern Shimanto belt) Northern Chichibu belt) (Figs. 5 and 6). On the surface, this boundary is corresponds to a north-dipping low-angle fault traditionally called represented by an essentially south-dipping low-angle fault called the Butsuzo Tectonic Line (BTL). To the north of Shimanto AC, non- Sasagatani Fault in central Shikoku (Isozaki and Maruyama, 1991; metamorphosed latest Jurassic to earliest Cretaceous AC unit (of the Kawato et al., 1991; Fig. 6A) between the high-P/T Sanbagawa meta-AC Sanbosan sub-belt in the Chichibu belt) occurs, structurally overlying and the non- to weakly metamorphosed Jurassic AC unit on the south. the former. The protoliths of the Sanbagawa meta-AC correspond in On the other hand, the boundary between the Sanbagawa meta-AC age to the Sanbosan AC unit, whereas the protoliths of the Shimanto and the newly recognized unit (tentatively called the Shimanto meta- meta-AC are identical to those of the N. Shimanto AC. According to AC hereafter in this article) is noteworthy. As shown in Fig. 6A, the this correlation based on OPs, the boundary fault between the surface trajectory of this boundary between the two blueschist units Sanbagawa meta-AC and the Shimanto meta-AC geologically corre- runs mostly parallel to topographic contour lines, particularly in sponds to the BTL. In fact, this subhorizontal contact was already eastern-central Shikoku, suggesting its low-angle nature. The Sanba- described between the N. Shimanto AC and an enigmatic meta-AC gawa meta-AC always occupies a structurally higher level, overlying unit (previously assigned to the “Sanbagawa high-P/T unit") in the the Shimanto meta-AC. In other words, the Sanbagawa meta-AC central Ki peninsula (Sasaki and Isozaki, 1992; Masago et al., 2005). Occurs as a nappe bounded by subhorizontal faults on both side, as The latter unit is properly identified as the Shimanto meta-AC. These Modernarc provenances Cretaceous-Tertiary arc provenanceregimeinJapan Yangtzecraton sandstone mid-Cretaceousarc Jurassicarc age (Ma) Permo-Triassic collision suture Sino-Korean craton EarlyPaleozoicarcs 600 Atlantic-type margin 600 1000 1500 2000 2500 3000Ma *520Ma:Beginningof ocear inicplatesubduction L.Sil 500 Pacific-type margin 400 1. Consumption of Paleozoic arc complexes .Tr Disappearanceof400-500Maarc .Jur M.Jur E.Crt *240-230 Ma:Collision of Yangtze and Sino-Korea E.Crt L.Crt 2. Consumption of earanceof150-250Maarc mid-Mesozoic arc complexes L.Crt Recnt 3. Consumption of Recnt post-Jurassic arc complexes * 20-15 Ma: Back-arc opening to form Japan Sea 100 200 300 400 500 600 1000 1500 2000 2500 3000Ma zircon age (Ma) Fig. 7. Age spectra of detrial zircon grains from the mid-Paleozoic to Mesozoic sandstones and Recent river sands in Japan, showing secular change in provenances that shed terrigenous clastics to Japan (modified from Nakama, in press). Note the 3 distinct stages in the over-5o0 million-year history of the Japanese Islands in terms of terrigenous clastics from granitic sources; i.e. before the Late Triassic (ca. 200 Ma), Jurassic to mid-Cretaceous (ca. 200-90 Ma), and after the Late Cretaceous (ca. 90 Ma). This suggests that a major provenance change occurred twice in the vicinity around Proto-Japan and the Japanese Islands. During the Paleozoic to Triassic, major source was Paleozoic arc granite that developed in the juvenile Japan that probably formed an intra-oceanic arc (proto-Japan arc complexes). Despite the physiological proximity, the terrigenous flux from the two major continental blocks (South and North China) was highly limited. The major re-organization in provenance regime around 2o0 Ma occurred in relation to the collision between South China and North China that caused the abundant production of terrigenous clastics and delivery to the arc-trench system along the Pacific rim.The second change around 9o Ma was likely induced by the establishment of the huge Cretaceous batholith belt in the Japan arc that formed a great barrier wall in front of the two old continental cratons to block the terrigenous flux. Y. Isozaki et al. / Gondwana Research 18 (2010) 82-105 93 results will promote a re-examination of older high-P/T units and (Nangrin and Yeongnam blocks) (e.g., Cluzel et al., 1990; Oh, 2006; their mutual boundaries in Japan. Fig. 1 bottom). According to its subhorizontal nature, the suture line is exposed twice on the surface in the peninsula, i.e., in the Imjingan belt 6.2. Hierarchy in geotectonic boundaries on the north and along the southern margin of the Ogcheon zone on the south. Further eastern extension, however, is unclear, probably Owing to the past and present unstable tectonic conditions, the due to the pinch-out of South China block per se. Japanese Islands are replete with faults. Among these, the MTL, Nonetheless, there is a possible candidate for its eastern continuity in Itoigawa-Shizuoka Tectonic Line (I-STL alias Fossa Magna), and Japan. The Higo belt in north-central Kyushu is characterized by Tanakura Tectonic Line (TTL) (Fig. 2) have been traditionally regarded medium pressure-type metamorphic rocks of ca. 230 Ma age (Osanai as major tectonic elements that constituted the backbone framework et al., 1999) that appears identical to the UHP metamorphism of the of the Japanese Islands. During the virgin field mapping over the terra suture. Ifthis is the case, the fragmentary distribution of the suture rocks incognito in Far East during the late 19th century, these clear solely in north-central Kyushu needs further explanation because there lineaments in Japan naturally became prime eye-catchers for the is no broad exposure of the North China basement in Japan. This is likely pioneer geologists who emphasized their geological significance. explained by the Miocene strike-slip dislocation of right-lateral offset However, these notions eventually became stigmas in mind for along the western margin of Japan Sea (as discussed below), and klippe- domestic geologists. Apparently, the prominent linear structures are like occurrence with respect to the erosion level on the surface (see inset all relatively new, mostly of Cenozoic origin, simply because they profile of Fig. 1). After all, the occurrence of this 230 Ma metamorphic were not overprinted/modified by later tectonics. Therefore, they S in de no iso sisns s u s could not be responsible in shaping the major backbone structures of intimately related to Cathaysia part of South China block. Japan in the past at least back to the Early Paleozoic. On the other hand, the significance of some narrow tectonized 6.2.2. Orogen-bounding fault zones accompanying serpentinite/high-grade gneiss in SW Japan The remarkable zonal arrangement of various deformed/meta- (Hida marginal, Nagato, Maizuru, and Kurosegawa belts) was morphosed rocks in SW Japan has motivated many geologists who emphasized during the 1950-1960s because these belts were tried to establish the geotectonic subdivision of the Japanese Islands regarded as large-scale crust-cutting fault zones that reached from with respect to orogenic context. These zones in SW Japan are the surface down to the upper mantle across the Mohorovicic composed mostly of the Paleozoic-Cenozoic AC/meta-AC units and a discontinuity (e.g., Ichikawa et al., 1956). In the early 1980s, these minor amount of passive margin elements on the Japan Sea side. To serpentinite-bearing belts were re-interpreted in slightly modified date, 12 distinct belts were discriminated, and accordingly 11 belt ways as deep-penetrating strike-slip zones (e.g., Taira et al., 1983) boundaries were drawn on the surface (Fig. 2). Despite considerable under the strong influence of the “"suspect terrane" concept from secondary modification in the Cenozoic, NE Japan essentially follows western North America (Jones et al., 1977; Coney et al., 1980). The the same subdivision as SW Japan. Most of the belt boundaries correspond to subhorizontal faults however, was eventually demonstrated as a klippe (isolated part of a between two neighboring and distinct AC/meta-AC units (Fig. 5) with spis yoq uo shun snon ne jo do uo ss (ddeu peuozouns unique OPS for each, and ages of AC units showing a sharp gap across without any root with modern deep crust/mantle (Isozaki and Itaya, each boundary with an overall downward younging polarity. In general, a peak orogenic episode is characterized by the exhumation of were likewise confirmed irrelevant to vertical mantle-reaching faults. a regional metamorphic belt and the formation of a granitic batholith In order to resolve the long-lasting confusion in evaluating geo- (Miyashiro, 1982). Judging from ages of high-P/T AC units and granitic logical significance of major structures in Japan, here we attempt to complexes in SW Japan,the Japanese Islands likely have experienced at categorize major tectonic boundaries/faults according to their geolog- least 5 distinct episodes of orogenic peak: episodes 1 to 5 at 450 Ma, ical and/or geotectonic context. As to the convergent plate margins in 340 Ma, 240 Ma, 120 Ma, and 60 Ma, respectively (Fig. 3). From the general, there is a clear hierarchy of tectonic boundaries in terms of orogenic viewpoint, five out of the 11 geotectonic boundaries are order of magnitude. Three distinct categories are discriminated from significant because they define the above-mentioned 5 major episodes. These are: 1) the base of the Oeyama meta-AC unit, 2) the bounding fault, and 3) secondary modifier. The first two form as a direct base of the Renge meta-AC unit, 3) the base of the Suo meta-AC result of the convergence between two major plates, whereas the third (+ Maizuru/Ultra-Tanba units) (= the Ishigaki-Kuga Tectonic Line; I- KTL; Isozaki and Nishimura, 1989), 4) the base ofthe Sanbagawa meta- involve either extensional, compressional and strike-slip movements AC unit (BTL), and 5) the base of the N. Shimanto AC (the Aki Tectonic within the overall convergent tectonic regime (e.g., Miyashiro, 1982). Line; ATL). Each of these boundaries essentially corresponds to ancient Wadati-Beniof plane along which the high-P/T meatmorphism took 6.2.1. Collision suture place (Fig. 8). Other belt boundaries between the AC units, e.g., the In the vicinity of Japan, there is only one major Phanerozoic suture base of Akiyoshi AC unit and that of the Mino-Tanba (-Chichibu) belt of continent-continent collision that is relevant to the evolution of the have relatively minor orogenic significance. It is noteworthy that all of Japanese Islands; i.e. the Dabieshan-Sulu suture between North China these boundaries between AC units are essentially subhorizontal, (Sino-Korea) and South China (Yangtze-Cataysia) blocks (Fig. 1). This therefore, they show sinuous trajectories on the surface (Fig. 2). suture zone running for more than 2oo0 km is characterized by the The boundary between the Oki belt and the Early Paleozoic meta- unique occurrence of 230 Ma ultrahigh-P/T (UHP) metamorphic belt ACs was named the Nagoto-Hida Marginal Tectonic Line (Ng-HmTL; s s o Isozaki and Itaya, 1991） because it has a profound significance of were all swept together during the termination of pre-existing oceans dividing the passive continental margin elements and the active (parts of Paleo-Tethys) before the final collision (e.g., Wang et al., margin ones (Fig. 8). Although this line has not yet been observed on 1989; Zhang et al., 2009). Due to the collision and successive sub- the surface, its location is expected along the Nagato Tectonic Zone in duction of South China beneath North China, the leading edge of the the western Honshu (Fig. 2). Some smaller fragments derived from North China structurally overlies South China. The eastern continu- the vicinity of this boundary occur also along the periphery of the Hida ation of this suture can be traced to the Korean peninsula, where the belt in central Honshu (Hida Marginal belt) where they constitute a South China basement appears in the middle (Gyeonggi Massif) as a chaotically mixed (mélange) unit formed by the Neogene emplace- tectonic window beneath the North China rocks on both sides ment of the Hida nappe from the north. The boundary between the 94 Y.Isozakietal./GondwanaResearch18(2010)82-105 Nagato-Hida marginal Tectonic Line (Ng-HmTL) South China Osayama Tectonic Line (new name) 55 Omi Tectonic Line (new name) Ishigaki-Kuga Tectonic Line (I-KTL) Middle Paleozoic meta-AC Late Paleozoic AC + meta-AC Butsuzo Tectonic Line (BTL) Jurassic AC + meta-AC Aki Tectonic Line (ATL) Early Cretaceous AC + meta-AC high-P/T meta-AC Paleogene-Recent AC Map view Profile OmiTL BTL Sanbagawa meta-AC 340Ma-WBplane 120-90 Ma WB plane (= Sanbosan AC) Osayama TL Shimanto meta-AC I-KTL 450Ma-WBplane 240Ma-WBplane (= N. Shimanto AC) Ng-HmTL Chichibu (= Mino-Tanba) AC 80-60 Ma WB plane Sanbosan AC fore-arcsediments ATL 40 Ma WB plane OkiB 0 km N. Shimanto AC S.ShimantoAC 10 20 Philippine Sea Plate 30 40 50 40 km Fig.8.Simplified anatomy of SW Japan in map view (above) and profile (below). Note the essentially subhorizontal stack of various AC and meta-AC units with a clear downward and oceanward younging polairty. The boundary faults are all flat-lying. Nagato-Hida marginal Tectonic Line (Ng-HmTL) represents a significant geotectonic boundary between the passive continental margin elements and active margin ones. The rest 5 major geotectonic boundaries in Japan all develop within the active margin elements (AC and meta-AC) and demarcate neighboring high-P/T meta-AC units; i.e., Osayama Tectonic Line (new name) at the bottom of 450 Ma meta-AC, Omi tectonic Line (new name) at the bottom of 340 Ma meta-AC, Ishigaki-Kuga Tectonic Line (I-KTL) at the bottom of 240 Ma meta-AC, Butsuzo Tectonic Line (BTL) at the bottom of 120 Ma meta-AC, and Aki Tectonic Line (ATL) at the bottom of 60 Ma meta-AC. Each of these boundaries corresponds roughly to ancient Wadati-Benioff plane along which the high-P/T metamorphism took place, although later tectonic erosion might have modified these boundaries. Nonetheless these boundaries divide the multiple episodes of mid-oceanic ridge subduction in the past. Y. Isozaki et al. / Gondwana Research 18 (2010) 82-105 S Hida belt and Oki belt has not been directly observed in the field on utilizing the mid-crustal detachment. This fault likely represents the land; nonetheless, this may have a unique geological connotation with origin of MTL (or paleo-MTL) as a low-angle south-vergent thrust that respect to the secondary modification as discussed below. The activated clearly prior to the Quaternary strike-slip fault (neo-MTL; Kurosegawa belt in SW Japan in fact represents a composite geologic Isozaki and Maruyama, 1991). This fore-arc contraction caused the entity that includes various pre-Jurassic ACs/meta-ACs. Its base- unnatural apparent juxtaposition of the Cretaceous granite batholith bounding fault forms a part of the I-KTL, enveloping the klippe (of the Ryoke-Sanyo belt) and nearly coeval high-P/T metamorphic (tectonic outlier) of the pre-Jurassic AC and meta-AC units. rocks (Sanbagawa meta-AC unit) side-by-side (Figs. 2, 6B); i.e., the paired metamorphic belts by Miyashiro (1961). A rapid denudation of 6.2.3. Microplate-relevant secondary modifier exhumed shallow crust (older ACs and granitic rocks) (Fig. 9) might accelerate the sedimentation at trench and also resultant growth of a arc, disturbing the major zonal arrangement of the AC/meta-AC units huge AC belt full of terrigenous grains derived directly from the in Japan; i.e. MTL, I-STL (Fossa magna), TTL, and syntaxis around the Cretaceous arc batholiths. Izu peninsula. These are not the primary structural features of the The activity of fore-arc sliver is another secondary modifier. Three orogenic edifice of the entire Japanese Islands but merely represent fore-arc slivers are recognized in Japan, the Kuril-E. Hokkaido sliver scars of younger, secondary tectonic modifications (Isozaki and (Kimura 1986), Nankai sliver (Isozaki, 1989), and Ryukyu sliver Maruyama, 1991). These include 3 types of deformation derived (Kuramoto and Konishi, 1989). Among these three, the Nankai sliver is from 3 distinct microplate movements (Miyashiro, 1982): 1) arc-arc best documented in terms of geology/geophysics. The Nankai sliver is collision suture (e.g., the Kozu-Matsuda Fault next to the syntaxis in central Honshu, West Hidaka Fault in central Hokkaido), 2) strike-slip local rift-related extensional basin on the east, and a seismically (partly transform) boundaries related to back-arc spreading (e.g., TTL, compressional domain on the west (Hasegawa et al., 2009, Hasegawa (-ou) s die-aon e jo Arepunoq (e pue ‘(eue esso et al., in press). It is noteworthy that the linearity of MTL becomes most As to arc-arc collision, Yamamoto et al. (2009) recently demon- prominent in the segment between the Ise Bay (on the south of strated that the subduction of modern arc complexes rarely cause Nagoya) and the Bungo strait (between Shikoku and Kyushu). This accretion of arc basement to a continental margin. Most of colliding arc complexes subduct smoothly beneath the leading edge of the Nankai sliver, proving that the right-lateral strike-slip movement of continental plate because the latter deforms plastically without the high-angle neo-MTL is driven by the oblique subduction of the leaving remarkable traces of indenters (e.g., Kyushu-Palau ridge). Philippine plate beneath SW Japan. On the basis of the seismic images The syntaxis structure in central Honshu is likely an exceptional of SW Japan, Sato et al. (2005) and Ito et al. (2009) emphasized the consequence of the orthogonal collision of the Izu-Bonin arc complex low-angle nature of the MTL at present; however, this fault probably that could accumulate additional crustal material in a particular represents the relatively old remnant of MTL,ie.,the contact between domain, where a small amount of arc crust has accreted since the the Sanbagawa meta-AC unit and the over-riding Ryoke (granite- Miocene, and a concave curvature was indented onto the zonal baked Jurassic AC) unit, formed during the Miocene (paleo-MTL). The arrangement of AC units (Fig. 2). Another rare example is the Maizuru downward continuity of the high-angle strike-slip fault (neo-MTL; intra-oceanic arc complex probably accreted in the Triassic. In Fig. 5) is likely overlooked in seismic interpretation. contrast, oblique subduction of arc makes the collision point migrate As mentioned above, these relatively new features are eye- laterally along the trench, without accumulating crustal material in a catching because of their clear lineaments on the surface. Nonetheless, particular domain, and is thus not effective in the accretion of arc their geological or orogenic significance is relatively minor, because complex. their contribution is limited to local modifications of the pre-existing A back-arc spreading occurred during ca. 20-15 Ma to open the main orogenic framework. In short, all of these secondary features Japan Sea and to isolate Japan as an island arc system. On the basis of may add some ornamentation to the subduction-related orogen but paleomagnetic data from the Tertiary rocks in Japan, Otofuji (1996) cannot construct the overall edifice with orogenic cores composed of proposed a “tavern-style" double-door opening model, by assuming the multiple regional metamorphic belts and granitic batholith belts. clockwise rotation of SW Japan and anticlockwise rotation of NE Japan. Therefore, the interpretations of faults in the Japanese orogenic belts The geometry of basins and seafloor magnetic stripe pattern of the Japan need to differentiate the relatively old, and thus more tectonically Sea, however, are not concordant with this model (Jolivert et al., 1994). Overprinted, orogens of the Precambrian. n s s s n o 6.3. Accretionary growth vs. tectonic erosion lozenge-shaped back-arc basin like the Japan Sea tectonically, it is inevitable to activate a pair of strike-slip fault systems on two sides of The geotectonic evolution of the Japanese Islands has been explained the basin. Along the eastern margin of the Japan Sea, an N-S trending to date as a simple, one-way process of continental growth toward the left-lateral strike-slip dislocation likely occurred to offset SW Japan Pacific ocean, regardless of either geosynclinal or plate tectonic relative to the south (Fig. 2). The TTL and associated parallel-running scenarios. These explanations were based on the observations of visible faults are candidates for the on-land expression of this eastern marginal zonal arrangement with oceanward younging polarity in Japan; fault system. On the other hand, the western margin of the Japan Sea is however, what have been long overlooked are some ancient units that demarcated by the N-S trending Ululun fault along the eastern coastline previously existed but are not seen at present. The recent provenance of the Korean peninsula. Judging from the geometry of local basins and analysis of detrital zircons has first imaged these “ghost" geologic units continental fragments within the Japan Sea (Fig. 2), the transform which once formed in Japan and have already disappeared without systems might play an important role in the opening processes of the evident traces. In addition, the establishment and termination of the back-arc basin. The I-STL (Fossa magna) may correspond to one of the sedimentary link between Japan and two major continental bocks remnants of Miocene transform systems that dissected Japan in a left- (North and south China) was also demonstrated. These new facts lateral sense. demand considerable modifications of the hitherto accepted geotec- Even during the back-arc spreading, a compressional tectonic tonic history of Japan, as discussed in the following sections. regime might appear in the fore-arc domain (Isozaki and Maruyama, 1991) because the leading edge of the southward moving crust of SW 6.3.1. Missing Paleozoic arc granites Japan was anchored by the subduction from the Pacific margin. The The mid-Paleozoic and Triassic sandstones in Japan contain abun- fore-arc domain of SW Japan likely suffered across-arc contraction by dant igneous zircon grains of the 520-400 Ma ages, suggesting that Y.Isozakiet al./Gondwana Research 18(2010)82-105 Arc-trench setting of Cretaceous SW Japan(ca.100-80Ma) 100-200km Volcanicfront Trench pre-Jurassic ACs JurassicAC SanbosanAC N. Shimanto AC T.L. -K BTL: P1 accretion anic crust ·ca.30km MOHO Kula-Plate Granitebatholith Sanbagawa (Ryoke + Sanyo) high-P/T meta-AC S. China margin. +fore-arcophiolite ca.100km Fore-arc contraction and Fore-arc shortening erosionalremoval of arc crust (ca. 20 Ma) 100-80Ma 100-80 Ma Voclanicfront Trench main erosional domain ca.50km→ Volcanic front Ryoke Trench granite pre-Cretaceous ACs erosion MioceneAC Back-arc P1 S.China Paleo-MTL sprending ATL accretion detachment fault PHS-Plate JurassicACs MOHO S. Shimanto AC -ca.100km Fig. 9. Simplified model of the Miocene back-arc spreading and fore-arc contraction for juxtaposing the granite batholith (Ryoke belt) and coeval high-P/T meta-AC unit (Sanbagawa belt) in SW Japan (modified from Isozaki, 1996). Above: The mid-Cretaceous arc-trench setting of the SW Japan segment in East Asia. The oceanic subduction from the Pacific side produced the AC belt (Northern Shimanto belt) next to trench, whereas a granite batholith belt (P1) (Ryoke belt) formed beneath the volcanic arc, in particular, immediately after the ridge subduction.The across-arc distance between trench and volcanic front is usually 100-200km.Below:the shortened Miocene SWJapan by the fore-arc contraction induced by the opening of the Japan Sea. By utilizing mid-crustal detachment, the upper crust of the arc (the Cretaceous batholith belt and associated pre-Cretaceous AC + meta-AC units) was horizontally transported oceanward. Consequently, the enigmatic occurrence of granite batholith unit over the coeval high-P/T meta-AC in western Shikoku was achieved. Note the location of P1 is ca.50 km from the trench. This fore-arc shortening likely accompanied severe erosion on surface to produce a huge amount of terrigenous clastics.Under the circumstances, an extensive AC belt (southern part of S. Shimanto belt) was formed, and also a large amount of clastics were possibly buried into deep mantle.In this case, the total volume of the Phanerozoic crust of Japan decreased, even though the oceanic subduction kept working. EarlyPaleozoicgraniteswere exposed/eroded extensively on the (Permian-Triassic) granites may also be correlated with the continental surface as the main provenance. No large igneous provinces (produced arc along South China. In the case of the 520-400 Ma (Cambrian- by mantle plume activity) were identified around the Phanerozoic Silurian) granites, however, there is no counterpart that remained in Japan. Therefore, these Paleozoic igneous zircons most likely came from East Asia, in particular in South China. Thus these Paleozoic granites may granitic batholiths of subduction-related magmatic rocks of arc origin. It have possibly formed in intra-oceanic arc settings. is generally difficult to identify whether these arcs represented The initiation of subduction-related magmatism around 500 Ma was continental arcs along the South China margin or intra-oceanic arcs off already pointed out before (Sakashima et al., 2003). However, the the continent, prior to the final amalgamation to Japan. The 190-160 Ma dominance of such Early Paleozoic zircon in the Paleozoic sandstone (Jurassic) and 110-90 Ma (Cretaceous) granites probably belonged to apparently contradicts the present limited occurrence of granites of the continental arc system that can be traced for more than 3000 km these ages as very small bodies of less than a few km across such as those along East Asia (Takahashi, 1983). Some parts of the 290-210 Ma in the Kurosegawa and Hitachi-Takanuki belts in SW Japan. A possible Y.Isozaki et al./Gondwana Research 18(2010) 82-105 N JapanSea SWJapan marginalsea Cenozoic OOkinawaIs. Ishigaki Is. 0 20°N East China Sea Pacific Ocean B NorthPala + 十30°N 135°E 140°E 500km Okinawa Is. Legend Okinawa Taiwan Trough +25N Post-Jurassic AC+meta-AC 120°E Ishiga Pre-Cretaceous AC+meta-AC 130°E Middle Mesozoic LatePaleozoic-EarlyMesozoic Continental fragments A B continentalfragments pre-CretaceousACs post-Jurassic ACs JapanSea Honshu OkiRidge Seto neo-MTL OkiTrough Ng-HmTLI-KTL BTL YamatoBasin InlandSea ugh SL SL Moho Lower Crust Mantle LC？ Philipine Sea Plate 50 Ol+Pxcumulate Mantle paleo-MTL 50 km DR km 600km 500 400 300 200 100 Fig. 10. The Phanerozoic AC (+ meta-AC) units in SW Japan and Ryukyus in map view and profile. A: The Phanerozoic AC and meta-AC units are divided into 3 major age groups; i.e. 1) Paleozoic and Triasc, 2) Jurassic, and 3) post-Jurassic (modified from Isozaki & Nishimura, 1989). The two major geotectonic boundaries are the Ishigaki-Kuga Tectonic Line (l- KTL) and the Butsuzo Tectonic Line (BTL).The width of older AC+ meta-AC belts is relatively smaller than those of the younger ones.B: The crustal profile of SW Japan (modified from Ito et al.,2009).Note the difference in thickness among the AC(+meta-AC)groups mentioned above.Thus in volume, the Paleozoic to Jurassic units are hardly recognized with respect to the huge ACs of the post-Jurassic age. This strongly suggests that older AC and meta-AC units are likely eroded away with time, even though the oceanic subduction kept working continuously, and that tectonic erosion is a possible process to make such an uneven age distribution of ancient AC units. explanation for this is that the Paleozoic granites once existed Maizuru, and Kurosegawa belts, and also on the southern vicinity of extensively in Japan, but they disappeared by Triassic. A similar MTL, in SW Japan. The discovery of the Early Paleozoic arc complex(es) situation can be identified for the 290-210 Ma and 190-160 Ma granites per se is significant in the paleogeographic reconstruction of proto- that show limited occurrence in modern Japan, such as in the Hida, Y.Isozakiet al./Gondwana Research 18(2010)82-105 Grenvilleorogen (1.0-1.3 Ga) Amz:Amazonia N.China Bal:Baltica ouog : 0 Kal:Kalahari 540 Ma Lau:Laurentia Rio:Rio de La Plata SF:Sao Francisco Gondwanaland Sib:Siberia W.Af:West Africa Japan W.Af Amz Pan-African Orogen SF Co Intra-oceanic arc Rio 1300-1000 Ma Amz:Amazonia Rodinia Afg:Afghanistan Au:Australia EAST Bal:Baltica E.Ant : East Antarctica GONDWANA Kal : Kalahari Ind :India Lau:Laurentia Rio:Rio de La Plata S.Chi : South China 60°S SF:Sao Francisco 30°S Sib:Siberia Japan W.Af:West Africa Fig. 11. Paleogeographic maps of the Meso- and Neoproterozoic supercontinents Rodinia (right) and Gondwanaland (left) (modified from Rino et al., 2008). Right: Rodinia was assembled through successive continental collisions that formed the Grenville orogen (dark blue area) during 1300-1000 Ma. In the Neoproterozoic, Rodinia broke up by the activation of mantle plumes, in particular, the large-scale Pacific superplume (the largest red circle). Left: Rifted continental fragments once surrounding Laurentia (North America) migrated away and re-assembled to form Gondwanaland on the other side of the globe by 540 Ma (Hoffman, 1991; Dalziel, 1992; Li, 1999). The original position of Japan within Rodinia was on the margin of South China, in particular, the Cathaysia part on the Laurentian side in a low-latitude domain. During the Early Paleozoic, South China was located somewhere close to the eastern part of Australia. disappeared by the Jurassic, without leaving any major vestiges, or volume of arc crust of proto-Japan became smaller than the total original shedding their clastics to the post-Triassic sediments in Japan. volume. Even in the case of the widely exposed Cretaceous batholith belt 6.3.2. Erasing arc batholith by tectonic erosion in SW Japan, the abundant coeval zircon grains in the Paleogene ACs In order for granitic batholiths to vanish within a long-lasting in the South Shimanto belt indicate that a huge portion of the subduction-related orogen, the following four steps are essential: batholith has been eroded, accreted, and presumably subducted in 1) extensive exposure of the batholiths on surface, 2) rapid erosion, part. These observations suggest that the continental growth along an active margin, i.e. the increase of juvenile crust, did not proceed accretion and/or burial into mantle. The abundant occurrence of unidirectionally, but was punctuated several times by severe Paleozoic zircons from the mid-Paleozoic to Triassic sandstones in shrinkage by the subduction of juvenile arc crust coupled with Japan per se proves that the first three steps actually have worked out deep burial into mantle. effectively to denude the Paleozoic arc batholiths. As to the step 4), as The Miocene back-arc spreading that opened the Japan Sea also long as these fragments/clasts of granite batholith remain as sediments drove the fore-arc contraction by activation of a subhorizontal mid- on surface within the same arc-trench system, the total volume of crustal detachment (paleo-MTL). Consequently, a huge amount of arc juvenile arc crust still remains the same. However, this is not the case for crust (pre-existing AC units, granite, and low-P/T meta-ACs) was the Paleozoic-Mesozoic Japan because the pre-Jurassic sandstones are eroded and transported toward the trench, thus abundant terrigenous of extremely small in volume. This indicates that most of the Paleozoic clastics filled the trench to build the Miocene AC (Fig. 9) and some arc batholiths were consumed almost completely and transported to parts of them were likely subducted into deeper levels. Judging from somewhere else. One of the possible processes to explain this situation is the seismic profile of modern SW Japan (Fig. 5), the deep subduction the subduction from trench into deep mantle. In other words, a part of of coarse-grained terrigenous clastics is not negligible in amount. This juvenile crust was lost from the fore-arc domain. Unless the subducted indicates that arc crust shrunk even during the active subduction sediments are totally underplated beneath the fore-arc domain, the net regime through tectonic erosion. Y.Isozakiet al./GondwanaResearch 18(2010)82-105 99 A.Early Ordovician(480 Ma) B.Early Devonian (400 Ma) GreatBritain Florida Caledonide Baltica S. Amer. urentia Africa Paleo-Asian lapetusocean ocean E Tarim Laurentia Ant. N. China Paleo-Pacific Aust ocean Indoch. Paleo-Pacific (Panthalassa) SouthChina proto-Japan ocean proto-Japan south China C. Early Paleozoic proto-Japan (520-400 Ma) proto-Japan arc proto-Japan sea (TTG,crust) 470Ma(oldest)sediments 520Ma(oldest) Paleo-PacificOcean granite trench SouthChina* 580 Ma (oldest) ophiolite 450 Ma (oldest) blueschist 520Ma(oldest) metasomatism Fig. 12. Paleogepgraphic maps around the Early-Middle Paleozoic proto-Japan. A: During the Ordovician, South China with proto-Japan was located still close to the eastern side of Australia (Maruyama et al, 1997); B: Some continental blocks, including South China and Indochina, started to move toward north during the Devonian; C: By 520 Ma (Early Cambrian), proto-Japan developed as an intra-oceanic arc system that featured subduction-related Tonalite-Tronjemite-Granodiorite (TTG) magmatic suite. The oldest metasomatism (in the Hida marginal belt) and the oldest high-P/T metamorphism (in the Oeyama belt; Tsujimori and Itaya, 1999) occurred within this system. The fore-arc sediments contain the youngest detrital zircon of 472 Ma (the Hida marginal belt; Nakama et al., in press-a,b). The oldest metagabbro from western Kyushu (ca. 580 Ma) likely represents the trapped oceanic crust next to the rifted continental margin.The oceanic domain between South China (Cathaysia)margin and proto-Japan,the proto-Japan Sea,likely had an ambient width enough to prevent the influx of continental detritus into the arc sediments. By checking and comparing the relative volume of ACs and meta- As demonstrated in Fig. 3, the accretionary growth during the ca. D ienpiu jo u a neu nude saoq i uedef u sa 500 million year subduction history of Japan occurred not continu- units becomes greater as its age gets younger, as illustrated in ously but intermittently. The previous explanation was that the Fig. 10. The Paleozoic ACs<Mesozoic ACs<Cenozoic ACs. In periods of non-accretion corresponded to those of highly oblique particular in the profile, we can confirm that the Cretaceous and subduction with strike-slip tectonic regime (e.g., Maruyama and younger AC units occupy almost all the fore-arc crust of SW Japan, Seno, 1986). However, it becomes clear that tectonic erosion was whereas the pre-Cretaceous ACs are almost invisible. These more effective than previously imagined. At present, it is more observations suggests that older AC/meta-AC units likely have reasonable to understand that accretionary growth likely occurred intermittently, and has alternated several times with tectonic subduction of sediments, i.e., tectonic erosion in subduction zone, erosion. This also suggests that the oceanward growth of continental because this mechanism offers the most probable solution for margin, in particular that of South China, did not occur continuously, erasing shallow-crustal material from the surface. Cenozoic- and the trench kept moving back and forth around more or less the Modern examples of tectonic erosion can be observed along the same position. In this regard, what can be better explained also is the Japan trench off NE Japan and also along the Chile trench where enigmatic occurrence of the chaotic Kuroseagwa belt with various older fore-arc crust has been actively eroded from the bottom (von mid-Paleozoic high-P/T metamorphic and granitic rocks that are Huene and Scholle, 1991). In nearly one half of the modern active exposed unnaturally too close to the modern trench. As originally trench, subduction of oceanic plate do not result in the formation of explained (Isozaki and Itaya, 1991), the nappe-like Kurosegawa rocks AC; instead the process tectonically erodes older fore-arc crustal material away and dumps them into deeper mantle. remained in the same position with respect to the wiggling trench. 100 Y.Isozakiet al./GondwanaResearch 18(2010)82-105 A. Latest Carboniferous (300 Ma) Siberia birthofsuper- B 280Ma) downweling N. China Siberia South China Japan unnamed Bureya plate Pacificocean 六 Laurentia Paleo -Tethys! N. China nian/Appalachian Dabie orogenidbelt r'seaway Farallon Tethya plate Paleo- m Carboniferous Austy Akiyoshi-Sawadan E.Ant nountchain South China Pacific ocean (Panthalassa) C. Midle Triassic (240 Ma) D. Late Triassic (210 Ma) E. Early Jurassic (180 Ma) Siberia Siberia Siberia FarallonPlate Bu Bu NorthChina FarallonPlate North China North China panmargin Carboniferous voshi-s nountchair South China SouthChir PermianAkasaka Kuzuu seamount chain IzanagiPlate Izanagi Plate Kuzuusear Fig.13.Paleogeographicmapsaround theLatePaleozoictoEarlyMe an (modified from Maruyama et al., 1997). Yellow: land; blue: sea; green: seamount and oceanic plateau; orange: granite batholith belt.A: Birth of the Asian super-downswelling (purple circle) at ca. 300 Ma (latest Carboniferous) and assembling multiple continental blocks, China, forming a suture along the Quinling-Dabie belt; D: By the Late Triassic, the suture developed into a mountain range that produced abundant clastics and exhumed ultrahigh- pressure (UHP) metamorphic belt; E: During the Jurassic, a wide accretionary belt developed, utilizing abundant terrigenous clastics from the collisional suture. Note the Permian intra-oceanic arc collided orthogonally against the Japan margin to accrete small amount of arc crust (the Maizuru belt). Thus, the ancient arc-trench system of proto-Japan was likely located whereas almost rare in South China. Accordingly, all the older sandstones more or less in the same position of the modern one, and the margin in Japan were previously believed to contain abundant older clastics ' Ainod se yn se umon a nou A udef jo derived from North China. From the conventional viewpoint, it is rather surprising to know that the Silurian sandstone in Japan is poor in 6.3.3.Precambrianconnection Proterozoic grains, and that the Proterozoic zircons are recognized only in It has long been recognized that the Jurassic conglomerate of the Mino- limited cases within the Cretaceous sandstones (Fig. 7). Tanba AC contains abundant clasts of Proterozoic granites and gneisses up Along the ancient rifted margin of South China, the tectonic to the 1.8 Ga age (Shibata and Adachi, 1974). The Proterozoic ages, 1500- turnover occurred from a passive to active regime by 520 Ma. A new 2500 Ma, are common in the North China basement (e.g., Jjin, 2002), n n s sn Fig. 14.Paleogeographic maps of Late Mesozoic to Cenozoic Japan (modified from Maruyama et al., 1997). A: During the Cretaceous, the Japan margin was transected twice by triple junctions of the trench-trench-ridge type.Two pairs of high-P/T meta-AC belt and granitic batholith belt were formed; B,C: During the Miocene, continental rifting started in East Asia to open up several marginal basins including Japan Sea, Bohai basin, Kuril basin, and South China Sea. The opening of Japan Sea ended around 15 Ma, after transporting oceanward the fore-arc crust of Japan to emplace coeval granitic batholith (Ryoke) belt over the high-P/T meta-AC (Sanbagawa) belt. Y. Isozaki et al./Gondwana Research 18(2010) 82-105 101 A.Mid-Cretaceous(90Ma) B. Miocene (25 Ma) Siberia Okhotsk Plate N.China IzanagiPlate Pacificplate S/China Proto-lzu-Bonin ambagawaBS eid Pacific Pacific KulaPlate Cretaceous Plate PhilippineSeaplate C. Late Miocene (15 Ma) NorthAmperica plate Eurasia plate ibas Pacific plate slab-melting Bohat 9.4 cm/yr paleo-MTL paleo-MT -Per Shikoku- Izu- -Bonin X PHSplate nu South China Sea Philippine Sea plate 102 Y. Isozaki et al. / Gondwana Research 18 (2010) 82-105 ocean boundary but probably within the oceanic domain of the proto- 7.2. Gondwana and proto-Japan (540-300 Ma) Pacific (or proto-Panthalassa), in the same manner as the modern Indo-Australia plate that is currently separating into two plates within The subduction of the Pacific seafloor started along almost all of the oceanic domain. The oldest oceanic fragment in Japan (580 Ma the continental margins around the Pacific (Panthalassa) by the Late metagabbro in the Nomo peninsula) may represent a piece of the Cambrian, ca. 500 Ma. Proto-Japan next to South China was located oceanic crust tectonically trapped between South China and the newly almost at the same corner of the Australian continent. The Japan established proto-Japan arc (Isozaki and Maruyama, 1991). This segment of South China on the Pacific side became incorporated into a trapped oceanic domain behind the Paleozoic arc was much wider subduction regime as an intra-oceanic arc system by ca. 520 Ma than previously imagined, because the Paleozoic fore-arc sandstones (Fig. 12A). At that time, proto-Japan was composed of two parts: a in proto-Japan are almost free from the Proterozoic detrital zircons. passive continental margin (the Oki belt) and an intra-oceanic arc Nonetheless, a minor amount of Proterozoic zircon grains were (the Oeyama belt). The latter was featured by arc batholith (Nakama detected in the Paleozoic and Triassic sandstones. This suggests that et al., in press-a,b), metasomatism (Kunugiza and Goto, in press), and the proto-Japan was located close to South China and the trapped high-P/T metamorphism of AC unit (Tsujimori and Itaya 1999). There was ample distance between the continental margin and the proto- drastic increase of the Paleo- to Mesoproterozoic zircon grains in the Japan arc, separated by a trapped ocean basin (Fig. 12C). Cretaceous time was probably due to the uplift of the Dabieshan-Sulu In the Early Devonian, at ca. 400 Ma, South China together with suture (Fig. 1) and associated extensive erosion of the North China proto-Japan arc was located still close to the Gondwanaland. The basement. By this time the proto-Japan Sea was terminated. After the collision between the Laurentia and Baltica formed Laurasia, closing rapid development of the Cretaceous batholith belt and the Miocene the remnant seaway of the Iapetus Ocean by successive subduction opening of the Japan Sea, the provenance link between North China (Caledonian orogenic belt colored in red) (Fig. 12B). South China and Japan was truncated by the formation of a large topographic moved to the north of Australia and became isolated from other barrier, such as a granitic mountain range and/or back-arc basin. continental blocks in the mid-Pacific. Proto-Japan grew into a more matured intra-oceanic arc-trench system with arc magmatism and high-P/T regional metamorphism (the Renge belt) (Nakama et al., in 6.3.4. Isolation of an island arc press-a,b; Tsujimori and Itaya, 1999). North China was also isolated The detrital zircon analysis for the Miocene formations in Japan on from the other continental blocks and belonged to a faunal province the Japan Sea side also clearly demonstrated that the modern Japanese distinct from South China. Islands became isolated abruptly from the main Eurasian continent at a particular time (ca. 16 Ma) as indicated by the total disappearance of 7.3. Pangea and the formation of Asia (ca. 300-200 Ma) Proterozoic grains (Nakama et al., in press-a,b), which is in accordance with other geological lines of evidence. During Carboniferous to Triassic, several continental blocks includ- ing South China, North China, Tarim, and Indochina moved northward 7. Paleogeography hole (Fig. 13A). In addition, the other continental blocks in modern Asia, A series of paleogeographic maps are presented here on the basis such as Siberia and Kazakhstan, were also swept into the same domain. This was caused by a large-scale super-downwelling in the mantle of the latest knowledge. These maps are basically modified from those (colored in purple) that developed beneath the domain corresponding by Maruyama et al. (1997) based on the new information synthesized in this paper. to the present Asia (Maruyama et al., 1997). On the other hand, the closure of the lapetus Ocean was completed along the Hercynian- Appalachian orogenic belts between Baltica and Laurentia, and so was 7.1. Rodinia breakup and the birth of Japan (ca.750-540 Ma) the Uralian seaway between Baltica and Kazakhstan/Siberia to make a large continent Laurasia in the northern hemisphere. Gondwana on the In the late Mesoproterozoic around 1.3 Ga, all major continental south and Laurasia on the north formed the supercontinent Pangea. blocks gathered together to form the supercontinent Rodinia, generat- In the Early Permian (around 280 Ma), North China and South ing several collisional sutures of the so-called Grenvillian age. Within China moved northward to merge with the Siberia, narrowing the Rodinia, a part corresponding to future South China was sandwiched seaways among these blocks. Proto-Japan was on the southeastern between Laurentia (present-day western side of N. America) and continental margin of South China and faced directly to the Pacific Australia/East Antarctica (Fig. 11 right). One of the important Ocean, or Panthalassa (Fig. 13B). The Carboniferous to Permian Grenvillian orogens currently observed in Asia is the Shibao orogen in Akiyoshi-Sawadani seamount chains (swarms) capped by shallow South China that binds the Cathaysia block (the southeastern half of the current South China craton) and the Yangtze block (the northwestern the Farallon plate subducted (Fig. 13B), and consequently some parts part) (Fig. 1). Cathaysia was once connected directly to Laurentia, of these were incorporated into the Permian accretionary complex whereas Yangtze to East Antarctica or Australia, respectively. The former (the Akiyoshi belt). An intra-oceanic arc system was colliding had a strong link to the birthplace of Japan (proto-Japan). orthogonally against Proto-Japan (Fig. 13C) and accreting a small By multiple superplume impingement and resultant rifting (Santosh amount of additional arc crust (the Maizuru belt). et al., 2009), Rodinia broke up during ca. 750-600 Ma and the proto- During the Triassic ca. 240-230 Ma, South China collided against North China probably from the Dabie promontory, closing the paleo- blocks were separated and migrated away, among which the major Tethyan seaway, forming the collision suture along the Quinling- pieces currently in the southern hemisphere collided on the other side of Dabieshan mountains in central China (Fig. 13C). The Paleo-Asian the globe to form the semi-supercontinent Gondwanaland by 540 Ma ocean (Mongolian seaway) between North China and Siberia also with a network of collsional sutures of Pan-African age (e.g., Hoffman, narrowed through the double-sided subduction system (Fig. 13D). 1991; Fig. 11 left). South China, with proto-Japan in its vicinity, was free Along the Pacific margin of South China, the active subduction of the from Gondwanaland as well as Laurentia, Baltica (northern Europe) and Farallon plate formed the Triassic ACs and meta-ACs (the Suo belt and Siberia (Li, 1999). The proto-Japan likely formed a segment of the Ultra-Tanba belt). passive continental margin of South China, in particular, the Cathaysia After 210 Ma, the Paleo-Asian ocean was mostly closed and a delta block on the Pacific side of the collisional sutures of the Pan-African age. Was formed in the eastern terminal part of the suture (Fig. 13D). The Y. Isozaki et al. / Gondwana Research 18 (2010) 82-105 103 Quinling-Dabieshan suture started to exhume the UHP metamorphic 1. The across-arc profile of the upper mantle beneath Japan confirmed rocks from mantle depths. The collision-induced surface erosion that the subduction of the Pacific plate continued for more than produced abundant terrigenous clastics that were transported to a 100 million years from a stable trench located at more or less the coastal delta to the northeast. This large amount of terrigenous same position for a long period. sediments were delivered to deep-sea trench along the active Pacific 2. By ground-breaking experiment and vibroseis analysis, the across- margin, and became the major building material for ACs. arc profile of the crust of Japan confirmed the subhorizontal stack of ACs in good accordance with the surface geology. 7.4. Growth of East Asia and Japan margin (200-25 Ma) 3. The protolith age of the fossil-poor metamorphosed ACs are precisely constrained by detrital zircon chronology which clarified Throughout the Late Triassic to Jurassic (Fig. 13E), owing to the that the traditional high-P/T Sanbagawa belt needs to be divided abundant supply of clastics from the suture, the accretion along the Japan into two parts; i.e., the Sanbagawa belt sensu stricto and the high-P/T margin constructed a huge belt of the Late Triassic-Early Cretaceous ACs metamorphosed AC of the N. Shimanto belt that represents the (the Mino-Tanba belt). Numerous fragments of Permian seamount youngest (ca. 80-60 Ma) blueschist unit in SW Japan. complexes of the Akasaka-Kuzuu chains capped by reef complexes were 4. The development of multiple Paleozoic to Cenozoic arc batholiths accreted. This tectono-sedimentary regime was fundamentally kept along in Japan was confirmed by detrital zircon analysis. This also the Japan margin through the Cretaceous until the Paleogene. Nonethe- suggests the potential of tectonic erosion to remove older arc less, two independent episodes of ridge subduction occurred back to back batholiths. during the mid-Cretaceous to punctuate the continuous formation of ACs; the first by the Izanagi-Kula ridge around 120-110 Ma and the sec- These new results require the re-evaluation of the existing concepts on ond by Kula-Pacific ridge around 70-60 Ma (Fig. 3). These two the geotectonic evolution of the Japanese Islands. The issues to be episodes contributed to build two pairs of high-P/T meta-AC belt on addressed in future include: 1) strict re-definition of all orogenic elements the ocean side and granite batholith belt on the continent side (Fig. 14A) and their mutual boundaries, 2) re-evaluation of the so-called accretion- ary growth of Japan, 3) re-consideration of paleogeographic reconstruc- 7.5.Establishment of the island arc system of Japan (ca.25 Ma-present) tion, and 4) formulating more-advanced model of the Miyashiro-type orogen that is formed through continuous oceanic subduction but The activity of the plume in sub-Asian mantle peaked to accelerate punctuated by episodic ridge-subduction. Further deployment of new the rifting the continental crust and to open major basins in East Asia, such as the Japan Sea, the Baikal basin and Bohai basin in northern China (Fig. 14B). Bimodal volcanism characterized the initial phase of on the ca. 500 million year old orogen in Japan. each rift-related basin formation. The Japanese islands became an island arc by the opening of the Acknowledgements back-arc basin during ca. 20-15 Ma (Fig. 14C). The back-arc of the This article is dedicated to the late Prof. Akiho Miyashiro for his strike-slip fault. The back-arc basin (the Japan Sea) reached its full great contributions in the studies on tectonics and also for his long- extent by 15 Ma, by transporting the fore-arc domain to the ocean term encouragement to younger geologists including the first author side. In contrast to the extension in the back-arc, the fore-arc likely of this article. S. Maruyama and M. Santosh provided constructive comments to the earlier version of the manuscript. A. Hasegawa and olith belt over the coeval high-P/T meta-AC belt by the subhorizontal D.P. Zhao kindly allowed the author to reproduce some of the figures. paleo-MTL. 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Terrane arrangement of Southwest Japan in view of the Paleozoic- Tsujimori, T., Itaya, T, 1999.lueschist-facies metamorphism during Paleozoic orogeny Mesozoic tectonics of East Asia.Association for Geological Collaboration in Japan in the southwestern Japan:Phengite K-Ar ages of blueschist-facies tectonic blocks Monograph 49, 145-155. in a serpentinite melange beneath Early Paleozoic Oeyama ophiolite.The Island Arc Yao, A., Matsuda, T, Isozaki, Y., 1980. Triassic and Jurassic radiolarians from the 8,190-205. Inuyama area, central Japan. Journal of Geosciences. Osaka City University 23, Tsutsumi, Y,Miyashita, A, Terada, K, Hidaka, H., 2009. SHRIMPU-Pb dating of detrital 135-154. zircons from the Sanbagawa Belt, Kanto Mountains, Japan: need to revise the Zhang, R.Y, Liou, J.G., Ernst, W.G., 2009. The Dabie-Sulu continental collision zone: a framework of the belt. Journal of Mineralogical and Petrological Sciences 104, comprehensive review. Gondwana Research 16,1-26. 12-24. Zhao, D.P., Maruyama, S., Omori, S., 2007. Mantle dynamics of western Pacific to East Uyeda, S., Miyashiro, A, 1974. Plate tectonics and Japanese islands: a synthesis. Asia: new insight from seismic tomography and mineral physics. Gondwana Geological Society of America Bulltin 85, 1159-1170. Research 11, 120-131. von Huene, R, Scholle, D.W., 1991. Observation at convergent margins concerning Zhao, D.P., Ohtani, E., 2009. Deep slab subduction and dehydration and their sediment subduction, subduction erosion, and the growth of continental crust. geodynamic consequences: evidence from seismology and mineral physics. Review of Geophysics 29,279-316. Gondwana Research 16, 401-413. Wadati, K., 1935. On the activity of deep-focus earthquakes in the Japan Island and neighbourhoods. Geophysical Magazine 8, 305-326.
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Isozaki (2010) New insight into a subduction-related orogen.txt
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Geochemical Journal, Vol. 22, pp. 183 to 193, 1988
Origin of some gases from
area inthe Takinoue geothermal
Japan
YASUHIRO Kiyosu' and YUTAKA YOSHIDA2
Department of Earth Sciences, School of Science, Nagoya University, Chikusa, Nagoya 4641
and Geothermal Development Division, Japan Metals and Chemical Co., Ltd. 24 Ukai,
Takizawa-mura, Iwate-gun, Iwate 021-012, Japan
(Received June 23, 1987; Accepted October 3, 1988)
Chemical and isotopic compositions were analyzed on geothermal steam from fumaroles and wells
in the Takinoue geothermal area. The distribution of deuterium and oxygen-18 in water samples sug
gests that the fumarole steam discharge is formed from the rising geothermal fluid through a single-step
steam separation process at temperatures of approximately 150-240°C. On the basis of N2/Ar, He/Ar
and C02/N2 ratios, and 813C values of carbon dioxide in steam, it is concluded that the gas components
dissolved in the Takinoue geothermal waters are mixtures of magmatic gases and meteoric waters.
INTRODUCTION
Geothermal fluids have been investigated to
evaluate temperatures in geothermal reservoirs
using solute or gas chemical geothermometers
(e.g., Fournier and Rowe, 1966; Fournier and
Truesdell, 1973; D'Amore and Panichi, 1980;
Arnorsson and Gunnlavgsson, 1985). Dissolved
constituents of geothermal fluids are derived
mainly from reservoirs and/or surrounding
rocks through water-rock interactions. Studies
of isotopic ratios suggest that the waters in most
hydothermal systems are mainly of meteoric
origin (Craig, 1963; Panichi and Gonfiantini,
1978; Giggenbach, 1978; Kiyosu, 1986).
However, the origin of gas components such as
carbon and sulfur species in geothermal fluids
has not yet been well discussed.
Yoshida (1984) found that the gases (N2, He
and Ar) of the Matsukawa vapor dominated
geothermal fields are mixtures of deep-seated
gases and atmospheric air dissolved in ground
water. Similarly, Kiyosu (1986) concluded that
small amounts of volcanic gases must have been
incorporated into the geothermal fluids from the
Onikobe area in Japan. However, the presenceof magmatic gases in geothermal systems has not
been clearly demonstrated.
The Takinoue hot water dominated type
geothermal area is one of the most active geother
mal systems in Japan (Fig. 1). Since May 1978, a
power plant of 50 MW has been in full operation
with geothermal steam from many production
wells. Based on chemical and isotopic steam com
positions, an attempt is made to relate the
fumarole steam to the underground geothermal
fluids. The origin of the various gas species in the
Takinoue geothermal field is discussed also.
GEOLOGICAL OUTLINE OF GEOTHERMAL FIELD
The Takinoue hot water type geothermal
area, Iwate Prefecture, is situated in the
Hachimantai volcanic group which belongs to
the Nasu volcanic zone. The geology of the
Takinoue geothermal field has been described by
Nakamura and Sumi (1981). The Neogene
sedimentary sequence, the Tamagawa welded
tuffs and Quaternary volcanic rocks are recogniz
ed in the Takinoue geothermal area. The
Neogene rocks are subdivided from lower to
upper in section into the Obonai formation,
183
184 Y. Kiyosu and Y. Yoshida
10
/ ,OD
C
114
POWER
STATION
NSea
Japan
Pacific
Ocean
I500m4:
5
2e12
k 1
700mn
909,
`809n
700
600m
600m
800m
Fig. 1. Location of sampling points in the Takinoue geothermal area,
and E: wells. A and B: SP-1; C: SP-2; D: KD-1; D and E.• SP-3.Japan. 1-14: fumaroles; A, B, C, D
characterized by dacite and dacitic tuff; the
Kunimitoge formation, consisting of black
shale, dacitic tuff, andesitic tuff and altered
andesite; the Takinoue-onsen formation,
characterized by alternating beds of argillaceous
rocks; and the Yamatsuda formation which is
composed of sandstone, siltstone and tuff. Weak
alkaline NaCI type hot springs and fumaroles are
widespread in the area. Acid-sulfate thermal
waters issue from mud pots in the surface altera
tion area. These waters have formed secondarily
by surface mixing of groundwater with S04 from
oxidation of H2S and condensed steam from a
boiling geothermal system at depth.
The rock alteration in this area consists of
montmorillonite and chlorite zones. Secondary
minerals such as pyrophyllite, kaolinite and
alunite occur locally around fumaroles and acid
hot springs (Kimbara et al., 1979).
Geothermal fluids are found in the highly
permeable zones created by fault and fracturesystems associated with folding in the Neogene
sedimentary sequence (Sato, 1982).
EXPERIMENTAL PROCEDURES
Steam samples were collected from the main
area of fumarole activity and wells from 1984 to
1985 by means modified from Ozawa's method
(1966). Figure 1 shows the location of wells (A,
B, C, D and E) and fumaroles (1~14) where
samples were collected. The fumarolic tempera
tures were measured with a mercury ther
mometer.
The analytical methods used to determine
water, carbon dioxide and hydrogen sulfide in
the samples are similar to those used by Ozawa
(1967). Other gases were analyzed by the gas
chromatographic method (Yoshida, 1984;
Kiyosu, 1985).
Isotopic analyses of water were carried out
by established methods (Bigeleisen et al., 1952;
Gases from Takinoue geothermal area, Japan 185
Friedman, 1953; Epstein and Mayeda, 1953).
The carbon dioxide in the samples was
precipitated as BaCO3, then reconverted to C02
using phosphoric acid.
Hydrogen, oxygen and carbon isotope ratios
were determined on a dual collector mass spec
trometer. The results are given in terms of a
value:
! X(%) _ (Rsample / Rstandard 1) X 103
where X indicates D, 180 or 13C; R refers to
D/H, 180/110 or '3C/'2C. The standard for
hydrogen and oxygen in all cases is SMOW and
that for carbon is PDB.
The overall reproducibility of samples was
±2.0%o for hydrogen and ±0.1%o for oxygen
and +0.2%o for carbon.
RESULTS AND DISCUSSION
Chemical composition of geothermal steam
The analytical results for gas constituents are
presented in Table 1. Total gas concentrations,
exclusive of H2O, are somewhat lower than those
found for Matsukawa (Yoshida, 1984) and
Onikobe (Kiyosu, 1986). The contents of gasesin the fumarole steam vary, whereas those in
well samples show a small variation.
The relative contents of the three main consti
tuents, C02, H2S and R-gas (alkali residual
gases) are given in Fig. 2. It is evident that the
data points fall into three areas. That is, steam
samples in the Takinoue fumaroles are divided
into three groups: 1) C02 type, 2) CO2-H2S type
and 3) H2S type. The C02 rich type steam is
found around the central zone of fumarole activi
ty, while the last two types of steam area associ
ated with the main fumarole activity. The H2S
poor fumarole steam samples suggest that
hydrogen sulfide is removed from the steam in
the upflow. It seems likely that oxidation of
hydrogen sulfide occurs when rising steam en
counters meteoric water.
On the contrary, the well steam corresponds
to the H2S rich type. The well steam samples sug
gest that the composition of gas in the reservoir
fluids is uniform. Conversely, the production
wells at the Matsukawa geothermal field are
divided into two groups: C02 type and H2S type
wells. CO2 rich wells are located in a zone of
weak alteration and other wells are located in the
zones where hydrous minerals such as mont
Table 1. Chemical composition of gas samples from the Takinoue geothermal area (vol. %)
SampleTemp.
°C H20 Gas C02 H2S N2 H2 CH4Ar
x 10-2He
x 10_4
T-1
T-2
T-3
T-4
T-5
T-6
T-7
T-8
T-9
T-10
T-11
T-12
T-13
T-14
KD-1
SP-1
SP-2
SP-397.0
98.0
99.0
99.0
98.0
98.0
98.0
97.0
98.6
98.0
98.099.93
99.90
99.96
99.96
99.95
99.92
99.94
99.94
99.93
99.96
99.98
99.67
99.62
99.94
99.98
99.97
99.960.07
0.10
0.04
0.04
0.05
0.08
0.06
0.06
0.07
0.04
0.02
0.33
0.38
0.06
0.02
0.03
0.0495.2
88.3
92.7
76.1
80.2
81.8
81.7
70.5
76.0
58.7
58.9
71.3
61.8
69.5
72.2
61.1
73.8
65.2 1.55
7.97
4.33
15.1
12.3
12.5
13.8
24.6
18.0
32.5
22.9
24.1
32.6
26.8
20.6
30.2
18.7
26.90.80
0.97
0.52
5.33
4.78
3.07
0.79
0.61
0.87
4.32
9.74
1.87
1.61
0.77
4.47
2.05
2.72
4.231.14
1.90
1.85
2.80
2.27
0.58
3.04
3.90
4.99
3.37
6.73
0.39
3.46
2.29
0.52
5.87
2.70
1.951.29
0.78
0.64
0.58
0.39
2.03
0.65
0.38
0.22
1.01
1.51
1.20
0.50
0.63
2.09
0.74
2.03
1.651.85
2.29
0.66
12.2
10.5
5.52
1.84
1.52
2.20
9.59
22.8
3.92
3.53
1.41
12.7
4.00
5.48
6.950.305
0.266
0.286
0.699
0.583
1.14
0.575
0.111
0.12
7.83
15.7
1.61
0.368
0.485
1.57
0.765
1.31
1.82
T.• Fumarole. KD, SP: Well.
186 Y. Kiyosu and Y. Yoshida
Nh oN4/0~
1Q/
•1
I
2•
6•,
• 7 5
4 II
9•
~11OSP-2
0
K D-1'12
*Be
14
S-3Ls
III
0 13
SP-1 •
1kP
4 ~osa
50 R' 25 0
Fig. 2. Relative proportions of CO2, H2S and R-gas (Residual gases) in the Takinoue geothermal steams.
R=N2+H2+CH4+Ar+He. 1-14: fumaroles; KD-1, SP-1; SP-2 and SP-3: wells. I: CO2 type; II: C02+H2S
type; III: H2S type.
morillonite, kaolinite and alunite are distributed
(Yoshida, 1984). This suggests that H2S concen
trations in the reservoir fluids increase with in
creasing contribution of hydrothermal altera
tion. The H2S concentration is probably controll
ed by the activity of sulfur in the water-rock in
teractions (Giggenbach, 1981; Kiyosu, 1987).
Although variation in fumarole gas
chemistry is observed in the Takinoue geother
mal field, its cause is not yet clear from the pre
sent study.Table 2. Isootope composition of water and cabon
dioxide from the Takinoue geothermal area (in %o)
Sample 6D 6190 613c cot
Isotopic compostion of steam
Isotopic data for deuterium and oxygen-18
are given in Table 2 and Fig . 3. Fumarolic steam
from Takinoue range from 71 .5 to 83.5%o
for aD and from -10.9 to -14.5%0 for x180 and T-1
T-2
T-3
T-4
T-5
T-6
T-7
T-8
T-9
T-10
T-11
T-12
M. W.-78.0
-83.5
-71.5
-72.6
-74.3
-75.4
-72.1
-80.9
-81.9
-72.5
-74.2
-69.0-12.2
-14.5
-10.9
-11.6
-11.8
-11.6
-11.4
-13.3
-12.3
-12.0
-11.7
-11.0-5.5
-5.8
-6.0
-7.6
-7.2
-7.2
-6.1
-7.4
-6.2
-7.5
-8.3
-6.8
M. W.: Kakkondagawa, river water . T. Fumarole.
Gases from Takinoue geothermal area, Japan 187
-40
60
0 O
r0
-80• Hot water
o Steam
o Local
meteoric water
steam from
diluted
deep water 180
~90/
0
2240*C
11 5 /Of C 2Well
200'C 7 2
0 O bo 0,
t0 steam from
c undiluted deep water
100r 100.c
jO of springs water after
steam loss
-15 -1018 -5 0 60 %0
Fig. 3. Isotopic diagram of steam and water formed from an original deep water of 240°C at various separa
tion temperatures or after dilution with local groundwater in the Takinoue area. Small open and closed circles
and tie lines show calculated isotopic compositions of steam and water separated at the indicated temperatures.
Open circles: data points for steam samples from fumaroles; Closed circles: NaCI type hot spring water samples
(Kiyosu, unpublished data); C-2: well sample (Matsubaya et al., 1985). M. W. L.: meteoric water line.
are consistently lighter than the surface water
(6D=-69.0%o and 6'80=-l1-0%o) Similar
results were observed in low temperature
fumaroles around some volcanoes in nor
theastern Japan (Kiyosu, 1983). This is inter
preted to be a result of liquid-vapor separation
by boiling, probably before or after the mixing
of ground water with geothermal fluids as de
scribed below. The well waters exhibit an oxygen
isotope shift of up to 2.7%o relative to waters on
the local meteoric water line having the same
deuterium concentration (Matsubaya et al.,
1985). This indicates that the reservoir water is
meteoric in origin..
The isotopic variation of geothermal fluids
during their rise to the surface results from deep
thermal waters that boil with decreasing pressure
and dilution by surface waters (Giggenbach and
Stewart, 1982).
According to Akeno (1978) the various solute
geothermometers in the Takinoue deep wellsgive reservoir temperatures of 220° to 260°C.
Therefore, we assume that the temperatures of
the original deep thermal waters before steam
separation or dilution, and meteoric waters are
240° and 10°C, respectively. The isotopic com
positions of thermal water and steam after boil
ing were calculated for the temperature range
1O0'-240'C using the following equations (1),
(2) and (3) by assuming that underground steam
separation from single phase thermal waters is a
single-step process.
asys+Csw(1 -Ys)=ao (1)
ys = (H0 HH) / (H5 Hw) (2)
aw=60+ 103(a-1)yy (3)
where 8w, 85 and 6. are the isotopic compositions
of thermal water and steam after boling and
original deep thermal water, respectively. ys in
dicates a steam fraction. Hw, HS and Ho are the
enthalpies of the thermal water and steam at the
188 Y. Kiyosu and Y. Yoshida
separation temperature and original deep ther
mal water, respectively. a is the equilibrium con
stant between liquid and steam waters. In Fig. 3,
variations in isotopic compositions of steam and
water after steam loss are given as a function of
temperature. Steam separation accompanying
adiabatic-expansion from 100° to 240°C leads to
a decrease in 8D and 6180 of steam and an in
crease in isotope compositions of thermal water.
Some data points for C02-142S and H2S type
fumarolic steam, and NaCI type hot spring water
lie close to the lines representing steam separa
tion from undiluted deep water at temperatures
ranging from 140° to 200°C. These separation
temperatures are consistent with the deuterium
(H2-H20) temperatures of 140°-220°C for
some Takinoue fumaroles obtained by Kiyosu
(1987).
On the other hand, the isotopic compositions
of the diluted thermal water produced by mixing
deep thermal water with local surface waters and
steams from its diluted water are obtained by
combining the following equations (4), (5) with
(3).
Joy,, + 8m(1 -yo) = aw (4)
yo =(Hw-Hm)/(Ho-Hm) (5)
where yo represents the fraction of deep thermal
water and Hm is the enthalpy of meteoric water.
8m indicates the isotopic composition of
meteoric water. Data points for C02 and C02
H2S type fumaroles occupy positions close to the
steam line separated from the original diluted
water.
The variation in isotopic composition of
fumarole steam is, therefore, likely to reflect the
effects of dilution and steam separation of ap
proximately 200°C from the original deep liquid
phase during their ascent to the surface.
Isotopic composition of carbon dioxide
Carbon isotope composition of CO2 in the
fumarole steams (Table 2) ranges from -5.5 to
-8.3%o with an average of -6.8%o, in good
agreement with those in values associated with
steam from wells ranging from 6.4 to 7.9%o(Yamamoto, 1981). This indicates that the car
bon dioxide of fumarolic and well samples at
Takinoue comes from a uniform carbon source.
These values are in accordance with those from
Matsukawa but lighter than the mean 3.4%o at
Onikobe (Kiyosu, 1986). The carbon isotopic
composition of C02 dissolved in surface water is
found to be 6.8%o which is close to the at
mospheric CO2 value.
Kiyosu (1984) found that volcanic carbon
dioxide as well as net-carbon discharges from
some active andesite volcanoes in northeastern
Japan have varying isotope ratios from 2 to
-10%o showing a trend that the VC value
becomes lighter as groundwater is increased.
This variation may indicate that mixing of car
bon dioxide derived from other carbon sources
with volcanic carbon occurred. On the other
hand, the magmatic carbon has an isotopic com
position in the range -6 to 8%o (e.g., Allard et
al., 1977; Deines, 1980) which is consistent with
the volcanic carbon. Because the carbon isotope
composition is mostly similar to that of
magmatic carbon, the geothermal carbon diox
ide in Takinoue does not appear to be significant
ly contaminated with other C02 sources such as
organic matter of VC = 25%o or marine
limestone (613C = 1 to 1 %o) during its ascent to
the surface from the reservoir.
Mixtures of magmatic and atmospheric car
bon could account for the Takinoue geothermal
C02 described below.
Origin of geothermal gases
N2 / Ar ratios of the fumarole and well gas
samples are plotted against the He/Ar ratios
together with those of the Matsukawa geother
mal system (Fig. 4). The data points from both
geothermal areas occupy positions along a line
that Kiyosu (1985) has evaluated in north
eastern Japan as having originated from
meteoric water toward magmatic gas. The mix
ing ratio of magmatic gases in the Takinoue and
Matsukawa geothermal fluids is -0.1 and
---1.0%0.
Yoshida (1984) pointed out that a small part
of these gas components at Matsukawa is deriv
Gases from Takinoue geothermal area, Japan 189
He/Ar
-2
10
10300
101
100
11•
12 14 •3
•SP.3
0001 ••SP.2
•6
SP. 1
KD•1 • 013
9104 0
5 Air
10 100 1000 2
Fig. 4. Relationship between N2/Ar and He/Ar ratios of geothermal gases. Open and closed circles indicate
the Matsukawa well samples (Yoshida, 1984) and Takinoue samples, respectively. KD-1, SP-1, SP-2 and SP-3:
wells; 1 14:fumaroles. Q: air dissolved in water at 10°C (N2l Ar=41, Hel Ar=1.5 x 10-4). Solid line shows
calculated simple mixing between magmatic gases (N2l Ar=4250, He/Ar=2.1) and air dissolved in water at
10°C (Kiyosu, 1985). The figures beside the curve indicate a fraction of the magmatic component in the mixing.N /Ar
ed from a deep seated source, not from the reser
voir. The same conclusion was obtained in the
Onikobe geothermal system (Kiyosu, 1986).
As shown in Fig. 5, although there is a cor
relation between C02/N2 and N2/Ar ratios, the
data points deviate from the mixing line which
represents the composition of magmatic gas mix
tures for all andesitic volcanoes with air dissolv
ed in groundwater (Kiyosu, 1985). One possibili
ty for this deviation is that an extraneous carbon
dioxide has been added into geothermal gases
with a constant CO2/N2 ratio, representing a mix
ture of magmatic gases and air dissolved in
groundwater. However, the carbon isotope of
geothermal CO2 would not point to such a possible soil or biogenic origin. The other possibility
is that even if these gases are completely dissolv
ed in the deep thermal water, vapor-liquid
separation of thermal waters during their ascent
to the surface leads to variations in the ratio of
gases present in a well discharge.
Assuming that the geothermal fluids
originate from the mixing of volcanic gases and
meteoric water at depth prior to boiling,
fumarolic steam corresponds to vapor phase
formed by adiabatic boiling of the deep fluids.
According to Giggenbach (1980), the discharge
gas content Xe,1 is represented by
XX,1=X1,;(1-y+yB,):'=X1,1Dt' (6)
190 Y. Kiyosu and Y. Yoshida
COIN 2 2
100
10 0o
0
1 0
7. 00 0
9 14
0 %
13SP100
16
SP-2
64
KD-1 • 110• SP-3
•11
0.010.1Q30.50.70.9
1 ~0 100 1000 2/Ar
Fig. 5. Plot of N2/Ar ratios versus C021N2 ratios of geothermal gases. Open and closed circles represent the
Matsukawa well steam (Yoshida, 1984) and Takinoue steams, respectively. KD-1, SP-1, SP-2 and SP-3: wells;
1-14: fumaroles. -4 : magmatic gases (Kiyosu, 1985). Solid line shows mixing between magmatic gases
(CO21 N2=270) and air dissolved in water at 10°C (CO21 N2=2.56 x 10-2). The figures beside the curve indicate a
fraction of the magmatic component in the mixing.
where X,,, indicates the mole fraction of liquid; y
is a fraction of vapor; B represents the coefficient
describing the distribution of gases between
vapor and liquid phase; D,=1-y+yB,. The plus
and minus signs refer to vapor gain or loss, re
spectively. Based on gas components, C1 and C2,
the gas ratio is defined by the following pressure
independent equation derived from equation (6),
(X1/X2)C=(X1/X2),(D1/D2)t1 (7)
At the same time, the geothermal fluids are
mixtures of magmatic gases and meteoric water
as represented in the following equation
(Kiyosu, 1985): (X1/X2)c=f(X1/X2)M+(1-f)(X1/X2)m, (8)
where M and m refer to magmatic gases and
meteoric water, respectively; f represents the frac
tion of magmatic gases. The variation in N2/Ar,
He / Ar and C02 / N2 ratios of geothermal fluids
expected from the boiling and mixing at 240°C
can be evaluated using equations (7) and (8) and
is shown in Figs. 6 and 7.
In summary, the variations in N2/ Ar and
He/Ar ratios are due to mixing, not boiling.
Conversely, the C02/N2 ratio varied due to the
boiling process. From Fig. 7, it is also found that
the gas components released from the Takinoue
geothermal system correspond to gases complete
Gases from Takinoue geothermal area, Japan 191
He/Ar
102
103h
O O OI g~ e
q ~ h J l
`°~lp-~QO J
~C Q C 001
S
f=0001
10 100 1000 2
Fig. 6. Relationship between N2/Ar and He/Ar ratios of vapor and liquid phases separated from an original
deep water of 240°C. Closed circle represents well samples. y is a fraction of vapor. f indicates a fraction of
magmatic gases. Q: air dissolved in water at 10°C.N /Ar
CO/N 2 2
100
100001~; 11 yy ~.
h ,
10
f_001 Ile
oQ x0~
0.1 1 0.3'0.50709
10 100 1000 2
Fig. 7. Relationship between C02/N2 and N2/Ar ratios of vapor and liquid phases separated from an original
deep water of 240°C. Closed circle represents well samples. y is a fraction of vapor. f indicates a fraction of
magmatic gases. -4: magmatic gases.N /Ar
192 Y. Kiyosu and Y. Yoshida
CO/ N2 and N2/Ar ratios.
1000
100
10Q3Q7W"Acknowledgments-The authors wish to thank Dr.
H. Shigeno of the Geological Survey of Japan for com
ments on the earlier version of this manuscript. This
study was supported in part by Japanese Grant-in-Aid
for Scientific Research, Nos, 58045062 and 59045067.
0..
.
.Q1
Q01
0
Q001
-10 -5 0
61CC02 aoa
Fig. 8. Plot of 813C values of CO2 versus C021N2
ratios. Dotted and closed circles indicate the well
(Yamamoto, 1981) and fumarole samples, respec
tively. -6-: magmatic gases dissolved in water at 240°C
which were calculated using data from Kiyosu (1984,
1985) (CO2/N2=2170, 613Cco1= -4.0%0).
ly dissolved in the deep liquid phase at depth,
and that the geothermal gases in this area are
derived from a mixture of magmatic gases and
meteoric water.
Figure 8 shows the plots of l 13C value for the
carbon dioxide versus C02/N2 ratio in liquid
phase from the Takinoue geothermal area. The
C02/N2 ratios of gases vary widely because of
the boiling process, but PC value of carbon
dioxide is rather uniform at approximately 6
to 8%o. Most of the data points lie close to the
mixing line between magmatic gases and CO2
dissolved in surface water, although the contribu
tion from the volcanic source appears to be
small. This finding is consistent with the results
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www.nature.com/scientificreports SCIENTIFIC REPORTS natureresearch OPEN ano spacinc along Sw pan arc caused bv ditference : in age of subducting lithosphere 8 z'@ oyuey enstey 's'zexo!yso yus!ous 'ebeuans !enqon 'e't!wnstel !yn!ysos Takumi Matsumoto* The SW Japan arc built by subduction of the Philippine Sea (PHS) plate exhibits uneven distribution of volcanoes: thirteen Quaternary composite volcanoes form in the western half of this arc, Kyushu Island, while only two in the eastern half, Chugoku district. Reconstruction of the PHS plate back to 14 Ma, together with examinations based on thermal structure models constrained by high- density heat flow data and a petrological model for dehydration reactions suggest that fluids are discharged actively at depths of 90-100 km in the hydrous layer at the top of the old (> 50 Ma), hence, cold lithosphere sinking beneath Kyushu Island. In contrast, the young (15-25 Ma) oceanic crust downgoing beneath Chugoku district releases fluids largely at shallower depths, i.e. beneath the non- volcanic forearc, to cause characteristic tectonic tremors and low-frequency earthquakes (LFEs) and be the source of specific brine springs. Much larger amounts of fluids supplied to the magma source region in the western Sw Japan arc could build more densely-distributed volcanoes. Subduction zone volcanoes tend to exhibit regular spacing along a volcanic arc, although the spacing of volcanoes within individual arcs is often variable from arc to arcl, 2. A broad positive correlation between the linear density of active volcanoes and the rate of plate convergence suggests that the faster subduction contributes to greater melt production in the mantle wedge3-5. Given that the fluids discharged from the subducted lithosphere drive magma generation, then a greater fluid flux is likely to enhance melt generation and may ultimately be linked to larger volcano numbers through increased formation rate of gravitational instabilities within the partially molten region in the mantle wedge4, 6-8. In addition to the rate of subduction, the slab temperature should also impinge on volcano distribution in arcs. The contrasting volcano density observed in the Japanese Archipelago may be attributed to the difference in slab temperature: the downgoing Pacific plate beneath NE Japan where active volcanoes are densely distributed is much older (ca. 200 my) and cooler than the PHS plate beneath SW Japan?. The abundant arc volcanism in NE Japan reflects partial melting in the overlying mantle wedge, triggered by active infiltration of slab-derived fluids, while most of the water in the warm PHS plate is driven off at shallow depths and is not available to trigger effective magma production in the mantle wedge beneath SW Japan?. This pioneering work estimated thermal structure of the PHS plate beneath the eastern part of SW Japan, the Chugoku district, where active volcanoes e o s n d s a (a ss canic arc exhibiting higher linear density of active volcanoes, in marked contrast to the Chugoku region of this arc, although both regions are underlain by the lithosphere of the PHS plate (Fig. 1). It should be here stressed, however, that the age of the PHS plate difrs contrastingly, 25-15 vs. ~ 50 my for lithospheres of the western these young and old lithospheres (Fig. 1). The diffrence in age hence temperature of the subducting slab could cause the contrasting volcano spacing along the SW Japan arc. Herein this hypothesis will be discussed based on the age of the subducted PHS plate inferred by reconstruction of plate motion, and calculation of temperature 1Kobe Ocean-Bottom Exploration Center, Kobe University, Kobe 658-0022, Japan. 2Department of Planetology, Kobe University, Kobe 657-8501, Japan. 3Research Center for Urban Safety and Security, Kobe University, Kobe 657-8501, Japan. *National Research Institute for Earth Science and Disaster Resilience, Tsukuba 305-0006, Japan.email: tatsumi@diamond.kobe-U.ac.jp SCIENTIFICREPORTS|(2020)10:15005 |https://doi.org/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ EUR Japan Sea Chugoku 35°N Kyushu 70mmly PAC B ShikokuBasin 30°N Amami- Sankaku Basin PHS 130°E 135°E 140°E Figure 1. Tectonic setting of the western part of the Japanese Archipelago, which is an orogenic belt activated by subduction of both the Pacific (PAC) and the Philippine Sea (PHS) plates beneath the Eurasian (EUR) plate building active volcanoes on the NE Japan and Izu-Bonin-Mariana (purple triangles) and the SW Japan arcs (white triangles), respectively. Quaternary volcanoes in the SW Japan arc are also shown by red circles. The thin and solid continuous lines denote depth contours to the top of the subducted PHS slab estimated based on the seismicity in the PHS slab and local-earthquake tomographyl4 and the broken lines show those estimated based on the teleseismic tomographyl4. Blue circles and a green belt indicate distributions of high Li/Cl brine springs and a LFE zone21, respectively. The thermal structures were estimated along the profiles A and B. distribution along the sinking PHS plate beneath the Chugoku and Kyushu regions, and the behavior of water during subduction of the PHS plate. Tectonic setting of the SW Japan arc. The PHS plate is being subducted beneath the Eurasian (or Amu- rian) plate along the Nankai Trough and the Ryukyu Trench at a rate of 40-70 mm/yearl3, in the northwest direction to form the SW Japan arc (Fig. 1). Earthquakes in the PHS slab take place actively down to a depth of ~ 150 km under Kyushu Island and ~ 80 km beneath the Chugoku region14 (Fig. 1). The PHS slab has been furtherinking aeismicallydowntoadepthof 400km516 Fi1)TectonictrmorsandFEswhichmay be caused by fluid activity asociated with dehydration of the downgoing slabi7, 18, have been identified beneath the non-volcanic forearc along the surface of the subducted PHS slab at depths of approximately 30-40 km with a belt-like along-arc distribution (Fig. 1). LFEs occur in high Vp/Vs areas, indicating the existence of fluid, near the up-dip end of the stable sliding region in western Shikoku19. Active arc volcanoes that are composed of lavas and volcaniclastics having the calc-alkaline signatures form 100-200 km above the top of sinking PHS plate in SW Japan, though the volcano density changes markedly between the Kyushu and the Chugoku segments (Fig. 1). In the non-volcanic forearc of the eastern half of this arc spring out characteristic deep-seated fluids referred to as the Arima-type brines (Fig. 1) possessing high Cl contents, high Li/Cl ratios, specific 818O-8D isotopic ratios, and high *He/*He ratios20,21 Such geochemical characteristics of these brines may be attributed to dehy- dration of the downgoing PHS oceanic crust21-24. SCIENTIFICREPORTS (2020)10:15005| https://doi.org/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ Kyushu Chugoku 120 (cwy) ues are higher than the equilibrium values at the experimen- the Happo spring (50-60 °C). Isotopic fractionation between gen- Spacing>20km3 * = 90 km 80 CH4 production from HzO with relatively small fractionation be- V/100km = 7.5 km3 e wni Vol 40 0 500 1000 1500 2000 2500 3000 Distance from the Southern Tip of Kyushu Island (km) Figure 2. The volume and location of Quaternary volcanoes along the SW Japan arc from Kaimon volcano at the southern tip of Kyushu Island29. Contrasting volcano spacing and volume of volcanics are observed for locations of such volcanoes are projected onto the volcanic front. In this diagram, volcanoes along the northern margin and on islands to the north and west of Kyushu Island are not included, because they are composed of alkaline basalts and may not be produced in association with subduction of the PHS plate. Some volcanoes in the SW Japan arc such as Aira and Aso volcanoes in Kyushu and Daisen in Chugoku erupted large amounts of felsic ignimbrites and/or tephra. These voluminous felsic magmas are not considered in this figure, since they may be produced by anatexis of the crust, not by differentiation of mantle-derived magmas30. The Japan Sea behind the SW and NE Japan arcs (Fig. 1) is a backarc basin created 30 to 15 Ma by rifting of the eastern margin of the Asian continent25. The opening of this backarc basin caused clockwise and counter- S a s (~50°C). The hydrogen isotopes of CH4-H2-H2O systems are as- be yielded in association with H2 generation and increasing pH separated from the paleo-Kyushu-Palau arc and migrated eastward, creating a new oceanic crust of the Shikoku opening and the clockwise rotation of the arc sliver, resulted in enforced subduction of the young (<15 my) Sankaku Basin behind the KPR on the PHS plate (Fig. 1) was born >48.7 Ma by backarc spreading within a e e e s go ed us a q pnpns q s pe i s re p es a rate of>65 mm/years (Fig. 1). It should be again noted that the age of the PHS plate changes greatly, 25-15 Vs. ~ 50 my, across the KPRi0-12 tion occurs at low temperatures owing to sluggish CH4-H2O iso- Volcano distribution in the SW Japan arc. Contrasting active volcano spacing is a characteristic in the SW Japan arc (Fig. 1). Identification of volcano spacing based solely on active Holocene volcanoes, however, are twofold. Firstly, although mafic melts can get transferred from source in the mantle wedge to surface rap- n s ( ) a o ou o-a topic re-equilibration during cooling The Happo methane and hy- the life span of arc composite volcanoes may be several hundreds of kilo years. In order to examine the linkage the differences in 813C-CH4 between seafloor and continental set- Japan, rather than that of active Holocene volcanoes, should be examined, because these volcanoes may be built by current motion of the PHS plate that have been constant since 3 Ma as described later. Figure 1 clearly exhibits that Quaternary volcanoes are much more densely built in the eastern half of this arc, although only two active topic exchange at 50°C. The 8(CH4-H2O) values from Happo #1 Figure 2 shows the along-arc distribution of Quaternary volcanoes and the volume of each volcano. It should be stressed in this diagram that most Quaternary volcanoes in the Chugoku region are small and form mono- such cases, H2O provides a source of hydrogen for methane for- with an average spacing of ~ 90 km, whereas only two in Chugoku (~ 500 km spacing). It may be thus confirmed that the contrasting volcano spacing in the SW Japan arc during Holocene has been continued from 2.6 Ma. There certainly is a gap in Holocene volcanism along the volcanic front in central to southern Kyushu Island (Fig. 1). Analyses based on receiver function33 suggested that this volcanic gap may be caused by migration of slab-derived fluids back to the forearc mantle wedge along the surface of the slab to form low-velocity, possibly serpentinized mantle. As indicated in Fig. 2, however, the volcanic gap may not be so clearly observed when SCIENTIFICREPORTSI > -160%o. The small 8(H2-H2Oaq) value would require high tem- https://doi.0rg/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ ★13-14 Ma Setouchi volcanism ★14 Ma plutonism activevolcano Q 35°N Chugoku Kyushu Trough 15-25Ma Nankai WarmPlate >50Ma Cool (>3 Plate Yaku-Shima Ma 30°N 56 130E 135°E Figure 3. Positions of the Kyushu-Palau Ridge (KPR) at present, 3, and 14 Ma arranged by subduction of the PHS plate shown by arrows. Red triangles, active volcanoes; green and yellow stars, Setouchi and near trench felsic volcano-plutonic complexes occurred 13-14 Ma. The cool and warm lithospheres bordered by the KPR underlay Kyushu and Chugoku at 3 Ma, respectively. reararc volcanos are included, suggesting the contribution of slab-derived fluids to arc magmatism even in this region. Further detailed analyses may be required for better understanding the cause of the volcanic gap along the volcanic front of central Kyushu. Figure 2 together with the above considerations then confirm that a larger number of volcanoes and the observation may intuitively lead to the conclusion that the older and cooler PHS plate to the west of the KPR the Chugoku region releases the water at shallow depths and cannot cause effective magma production in the mantle wedge. However, this simple mechanism could not be applied, since the boundary between the older and younger PHS plate, i.e., the KPR is currently located beneath the southern part of Kyushu Island (Fig. 1). Paleo-position of Kyushu-Palau Ridge (KPR): contrasting age of subducting Philippine sea plate. The northern tip of the KPR, a remnant conjugate arc of the active Izu-Bonin-Mariana arc system, is located presently at the junction of the Nankai Trough and the Ryukyu Trench and is sinking beneath Kyushu Island (Figs. 1 and 3). The KPR plays a key role in the volcano-tectonic evolution of the SW Japan arc, as this forms a boundary between a younger (<25 Ma) and an older (>50 Ma) oceanic lithosphere and is composed of buoyant arc crust with the midle crust exhibiting seismic velocity similar to that of the bulk continental crust34. Although it has been accepted generally that the subduction direction of the PHS plate changed from NNW to NW35, the timing of this change has been controversial. Geological and structural evolution of strata deposited in the forearc basin of at the eastern margin of the PHS subduction system, however, has led to the conclusion SCIENTIFICREPORTS (2020)10:15005| https://doi.org/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ that it took place at 3 Ma and caused the stress change both in the NE and SW Japan arcs36.37. If so, then the paleo-position of the KPR at 3 Ma could be reconstructed based on the current motion of the PHS plate (Fig. 3); the KPR was situated beneath the northeastern edge of Kyushu Island at 3 Ma. It is thus suggested that a> 50 Ma cool plate has distributed beneath the Kyushu region, whereas a warm lithosphere of the Shikoku Basin beneath the Chugoku region in SW Japan. Characteristic volcanic rocks including mantle-derived high-Mg andesites erupted sporadically at 13-14 Ma and formed the Setouchi volcanic belt in the present forearc (Fig. 3), extending for ~ 600 km with five major volcanic regions38. Synchronously with this magmatism, formed felsic volcano-plutonic complexes at 14 Ma39. 40 in the near-trench region of SW Japan (Fig. 3). If the dip angle of subduction has been unchanged for the last 14 my, the slab depth beneath these forearc or near trench magmatic belts would have been<50 km, much shallower than that beneath most arc volcanic chains (110-170 km)*. Magma generation above such a shallow slab would require some additional conditions such as unusually high temperatures in the sinking lithosphere with subduction of a newly-born lithosphere of the Shikoku Basin enforced by southward migration of the SW Japan arc sliver in association with opening of the Japan Sea41. If so, then the distribution of 13-14 Ma forearc magmatism in SW Japan could provide a constraint on the location of the KPR at that time; to the south of Yaku- Shima Island (Fig. 3). The rate of northward subduction of the PHS from 14 to 3 Ma can be then calculated as 77 mm/years, almost identical to the previous estimation42. Thermal structure and dehydration of the subducted PHS plate: contributions to arc mag- matism. To understand migration of aqueous fluids associated with subduction of the young vs. old PHS plates and its role in causing contrasting volcano spacing in the Kyushu and Chugoku regions of the SW Japan arc, the thermal structures beneath two regions were estimated by 2D thermal structure models (Figs. S1, S4 and S6). Two end-member models are here constructed: One is a simple model (MODEL I) with constant slab model (MODEL II) that takes into account of the history of the subducted Philippine Sea plate and fits heat flow data best (Figs. S2, S3 and S5). In comparison to 2D models for these regions reported so far43-47, our thermal modeling has the following two advantages: (1) various heat sources were considered in the energy equation and (2) highly densely distributed heat flow data were used to constrain thermal structures and to estimate optimal values of model parameters in MODEL I (Fig. S5). Details of modeling are described in “Methods” and Supple- mentary Information (Sl). Although 3D thermal modeling may provide new insights into the thermal evolution several hundreds of different values of model parameters with high spatial resolution must be tested to obtain high-resolution thermal structure suitable for examining the behavior of water in the subducted slab and the mantle wedge. The calculated pressure-temperature (P-T) paths near to the surface of the PHS plate along the two profiles by MODEL I are shown in Fig. S4, together with HO contents in the subducted oceanic crust and the downgo- ing peridotite under HO-excess conditions calculated by Perple_X50 for a Shikoku Basin basalt and a peridotite (Table S2). It is indicated that temperatures in the two regions increase remarkably at a depth of ~ 40 km because this model does not incorporate a possible decoupling depth. As discussed later, an important point is that a o np sou (sn ad e n st (o) g ad e aa difference between them. In MODEL II, we examined two possibly most preferable models among several hundreds of models with different values of model parameters in terms of the least square sense of the observed heat flow data along profiles A and B (Fig. S5b,c): One is the cold forearc model (MODEL II-1) incorporating decoupling to depths model (MODEL II-2), which exhibits a remarkable temperature increase around a depth of 40 km along the plate interface as obtained by MODEL-I (Fig. S4). However, it is difficult to identify which model is better, because the difference in heat flow calculated for these two models is small and these heat flow values are largely consistent with the observed heat flow data. Further considerations on these two models are presented in the section “Pos- sible decoupling depths" in SI. The estimated most suitable values of the model parameters for the cold forearc model along the two profiles are tabulated in Table S1. Although three models, MODELs I, I-1, and II-2, provide different P-T profles along the subducting crust profile A mostly due to the age difference in the two regions. The behavior of HzO along with subduction of the PHS plate shall thus be examined based on the positively close-to-reality end-member model (MODEL II-1). It should be stressed, on the other hand, that discussions based on other models (MODELs I and II-2) reach to the conclusion on the cause ofthe contrasting volcano spacing in the Kyushu and Chugoku regions exactly the same as that based on the MODEL II-1. The present result for the Chugoku profle confirms the previous suggestion9 43 that the oceanic crust sink- ing beneath this region is warm and most of the water in the oceanic crust is driven off at shallow depths not to trigger partial melting of the mantle wedge directly and may further provide insights into migration of fluids and its role in characteristic fluid-related activities in this subduction zone. One is the occurrence of tectonic tremors and LFEs17,18,43,5, 5 taking place at ~ 30 km depths (Fig 1). The major dehydration reaction in the subducted PHS oceanic crust corresponding to the transition to amphibolite facies takes place 20-40 km depths (Figs. 4 and 5) and triggers characteristic tectonic tremors and LFEs. The other distinctive subduction-related fluid activity in this region is the occurrence of deep-seated fluids exhibiting characteristic chemistry and often high temperature. Water discharged from the PHS crust migrates upwards to form hydrous peridotites at the SCIENTIFIC REPORTS| I S005T:01 (020z) https://doi.org/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ Kyushu Chugoku 160- 160- Oceanic Crust Oceanic Crust 140 140- bpns 120- 120- VsEC qeis 100 VF 100- Depth (km) LwsBS 80 80 DryEc 60- 60 BS ZoEC 40- AmpEC 40 GR 20- AMP 20 300 600 006 1200 300 600 006 1200 160 160- Mantle Mantle 140 140- 5kmbelow 2.5 km above 120 120 100 100- Depth (km) amphibole 80- 80 60 60 40 40 20 20 0- 0 300 600 900 1200 300 600 900 1200 Temperature (°C) Temperature (°C) 3 4 5 >6 H2O content (wt.%) Figure 4. The calculated temperature distributions at the surface of the PHS plate along the profiles A and B passing through Kyushu and Chugoku districts, respectively for MODEL II-1 (Fig. 1). The metamorphic facies for the basaltic system, the stability limits of hydrous phases in the peridotite system, i.e., serpentine, chlorite, and amphibole, and H2O contents are also shown. GS green schist, EA epidote amphibolite, BS blue schist, AMP amphibolite, GR granulite, AmpEC amphibole eclogite, ZoEC zoisite eclogite, LwsEC lawsonite eclogite, DryEC dry eclogite. SCIENTIFICREPORTS (2020)10:15005| https://doi.0rg/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ Dense Volcanism (a) Kyushu 0 20 40 Higher 60 water flux 80 Eclogitetransition Chlorite+apmphibole in the subducted crust D dehydration 100 in the downdragged hydrous layer 120 140 一 350 300 250 200 150 100 50 Distance from the trench axis (km) Sparse Brine Volcanism (b) Chugoku Spring 0 Serpentinedehydration / in the downdragged 20- layer 40 Lower + Chlorite dehydration (km) water 60 flux in the downdragged Amphibolite transition layer in the subducted crust pth 80 Eclogitetransition in the subducted crust D 100 Amphiboledehydration in the downdragged hydrous layer 120 140 400 350 300 250 200 150 100 Distance from the trench axis (km) 0 2 3 4 5 6 H2O content (wt%) Figure 5. Dehydration and fluid release taking place along the subducted PHS plate beneath Kyushu (a) and Chugoku (b) for MODEL II-1. Higher water flux beneath Kyushu contributes to dense volcanism, while lower water flux caused by major dehydration reactions at a depth of ~ 50 km beneath Shikoku could not produce magmas actively to build a sparely-distributed volcanoes. base of the forearc mantle wedge, in which serpentine, chlorite and pargasitic amphibole may crystallize as major hydrous phases (Figs. 4 and 5). The hydrous peridotites are likely to be dragged downwards on the slab as a consequence of subduction of a rigid oceanic lithosphere into the viscous mantle, to supply aqueous fluids to the overlying dry mantle wedge°. Figure 4 also demonstrates that dehydration of serpentine and chlorite at the base of the hydrous peridotite layer ocur, i.e., immediately above the slab surface, at ~ 40 km to release large amounts of water, which would be the source of characteristic deep-seated fluids referred to as the Arima-type brines (Figs. 1 and 5) as advocated geochemically1-23, 41 Amphibole is then a hydrous phase in the down-dragged hydrous layer on the PHS plate that could trans- p s s e asa r p e e aso p sia dap o sa corresponds to the top of the PHS plate immediately beneath the active Sanbe volcano and other Quaternary volcanoes (Fig. 1). The Sanbe volcano is known by the occurrence of adakites exhibiting anomalously high Sr/Y ratios, leading to the conclusion that partial melting of the eclogitic PHS crust to cause magmatism9, 53. However, SCIENTIFIC REPORTS| | S005T:0T (020z) https://doi.org/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ geochemical examination of adakites from this volcano and the surrounding suggests that these adakitic mag- mas are produced by melting of the lower crust, not of the subducted oceanic crust, in the presence of garnet, plagioclase, and amphibole54-56. It should be thus stressed that the amount of water supplied through amphibole dehydration in the down-dragged hydrous layer to cause arc magma production is much smaller than that in the original hydrous peridotites including serpentine and chlorite. older (Fig. 3) and hence much cooler (Fig. 4). Providing that near the surface of the oceanic crust is significantly hydrated, then the subducting slab may largely dehydrate at 50 ~ 80 km depths (Fig. 4). In contrast, the hydrous layer at the base of the mantle wedge, i.e., immediately above the slab surface, enables to transport of a large amount of water to deep levels; chlorite and amphibole decompose to release water at depths of 100 km (Fig. 4), than a warmer thermal regime such as the Chugoku profile (Fig. 5). It has been observed that hydration of mantle portion of the oceanic plate may occur at least in some outer- rise regions through bending-related faulting prior to subduction57, 58. The Pacific plate being subducted at the Japan Trench, for example, the presence of ~ 2 wt% HzO in the uppermost mantle immediately below the Moho could account for the observed seismic velocity reduction58. If this is the case for the subducting PHS plate, then as serpentine or chlorite decomposes to release H,O beneath the volcanic chains (mantle temperature profiles at 5 km below slab surface in Fig. 4). Even if this is the case, then the contribution of serpentine dehydration to magma generation may be inferred beneath the Kyushu, not Chugoku region. The PHS plate is sinking normally along the Ryukyu Trench at the rate of>63 mm/years, whereas obliquely along the Nankai Trough at the rate of 61 mm/years with a substantial trench-normal component of~ 55 mm/ years. The effective rate of subduction of the PHS plate is higher in the Kyushu than Chugoku regions. This may enhance the contrasting volcano distribution along the SW Japan arc, as higher rates of subduction tends to cause higher rates of magma production in the mantle wedge3- Concluding remarks Quaternary volcanoes are distributed much more densely in the Kyushu than in the Chugoku segment along the SW Japan arc, although the PHS plate is currently being subducted beneath this arc. Tectonic reconstruction of the PHS plate that changed its direction of motion from NNE to NW at 3 Ma suggests an older (>50 Ma) por- tion of the PHS oceanic crust with high dip angle has downgone beneath the Kyushu region, whereas the young (25-15 Ma) lithosphere of the Shikoku Basin with low dip angle in the forearc region of the Chugoku segment of the SW Japan arc. Geothermal calculations of the temperature distribution along the subducting PHS plate with different ages, together with petrological constraints on dehydration reactions taking place within both the downgoing crust and the overlying mantle wedge, demonstrate that much larger amounts of fluids are supplied to the magma source region beneath the Kyushu than the Chugoku regions, causing much higher density in volcano distribution in Kyushu. Water that are released from the young PHS plate beneath the forearc of the Chugoku region may cause characteristic tectonic tremors and LFEs, and be the source of brain springs. Methods The calculation of 2D thermal structures in this study follows the previous models47, 51, 52. The momentum and energy equations were solved as a coupled problem, using the finite difference method. The model is a time- dependent, and considered possible heat sources such as viscous dissipation, adiabatic compression, frictional heating on the plate interface and temperature change caused by erosion and sedimentation during the Quater- creep59, and the density depends solely on the temperature. of 800 km and a depth of 400 km. Both the upper and lower crusts were set as conductive layers with respective thickness of 16 km. The accretionary prism was also incorporated into the model as a conductive layer. The thick- ness of the PHS plate at the Nankai Trough at the right model boundary is given based on the equation related to the half-space cooling with adiabatic compression at depths deeper than 50 km. As the boundary condition for flow fields, the normal stress is set to zero for the left, right, and bottom boundaries. As the boundary condition for temperature field, the model surface is set to O °C. Adiabatic conditions are assumed for the left and bottom prescribed guide62 whose length gradually extends from the right boundary from 14 Ma. Grid sizes for stream functions and temperatures are 4×4 km and 2 × 2 km, respectively, and the stream function is evaluated at the same grid spacing as the temperature field via the third-order Spline interpolation. Remeshing with 1 km for the mantle wedge corner, where intense flow is expected to occur, is performed at each time step63 64. For a simple model (MODEL I), we gave constant age of 50 my and constant velocity of 64 mm/years for the subducting PHS plate along profile A, whereas those of 17 my and 44 mm/years along profile B throughout the calculated period of 14 my (Figs. S2 and S3). We did not use heat flow data, and a decoupling depth is not incorporated into the model. On the other hand, for a rather complex and positively close-to-reality model (MODELs II-1 and 1I-2), the depth and age dependent temperature distribution determined by the plate cooling model RT165 is imposed at the right boundary. Time-dependent age and subduction velocity along profiles A and B were given, following the assumed subduction history (Figs. 1, 4, S2 and S3). For MODELs II-1 and II-2, we also used heat flow data from high-quality high-density Hi-net borehole and BSRs, which have not been used except for studies of our SCIENTIFICREPORTS| I S005T:01 (020z) https://doi.org/10.1038/s41598-020-72173-6 www.nature.com/scientificreports/ density heat flow data (Fig. S5a) to constrain the thermal structures along the profiles passing through Kyushu and Chugoku regions. This enables us to estimate thermal structures with high spatial resolution from shallow to deeper portions in association with subduction of the PHS plate. We correctly picked up only data along the two profiles within one-sided width of 30 km (Fig. S5b,c), Tables S2 and S4), and estimated optimal thermal models in which the calculated heat flow fits best with the observed values by least square method. It should be noted that spatial distributions of the observed densely-distributed heat flow along the two profiles obtained in this study are rather different from those of previous studies44; Shorter wavelength patterns can be identified, which should be explained by introducing heat sources such as temperature change caused by erosion and sedimenta- tion during the Quaternary period46. To better reproduce the observed heat flow data along the two profiles, pore of the PHS plate at the Nankai Trough, age discontinuity passing through the KPR, depth range and thickness of a low-viscosity layer attached on the plate interface, and its viscosity contrast against the surrounding region are assumed to be unknown free parameters (Table S1). Then, we performed grid search for several hundreds of different values of such free parameters for the respective profiles. Other details of the thermal modeling are described elsewhere47,51, 52 Data availability All data generated and analyzed in this study are included in main text or Supplementary Information. Received: 7 June 2019; Accepted: 16 August 2020 Published online: 14 September 2020 References 1. Shimozuru, D. & Kubo, N. Volcano spacing and subduction. In Arc Volcanism: Physics and Tectonics (eds Shimozuru, D. & Yokoy- ama, I.) 141-151 (Terra Publisher, New York, 1983). 2. Gill,J. 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A community benchmark for subduction zone modeling. Phys. Earth Planet. Int. 171, 187-197 (2008). 65. Grose, C. Properties of oceanic lithosphere: revised plate cooling model predictions. Earth Planet. Sci. Lett. 333-334, 250 (2012). 66.Wessl P, Smith,W.H E, Scharr, R, Luis, J F &Wobbe, F Generic mapping tools: imroved version released. EOS Tran AGU 94, 409-410 (2013). Acknowledgements We thank Harue Masuda for constructive discussions. This work was partly supported by JSPS KAKENHI grant were produced by using Generic Mapping Tool, GMT, version 5.3.366. Authorcontributions Y.T. and S.Y. conceived this study. S.Y, N. S. and T. M. conducted thermal modeling and Y. T. and K. K. is respon- sible for tectonic and petrological examinations. SCIENTIFIC REPORTS| (2020)10:15005| https://doi.org/10.1038/s41598-020-72173-6 10 www.nature.com/scientificreports/ Competing interests The authors declare no competing interests. Additionalinformation Supplementary information is available for this paper at https://doi.org/10.1038/s41598-020-72173-6. Correspondence and requests for materials should be addressed to Y.T. Reprints and permissions information is available at www.nature.com/reprints. Publisher's note Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affliations. License, which permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons license, and indicate if changes were made. The images or other third party material in this article are included in the article's Creative Commons license, unless indicated otherwise in a credit line to the material. 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Earth and Planetary Science Letters 554 (2021) 116656 Contents lists available at ScienceDirect EARTH Earth and Planetary Science Letters ELSEVIER www.elsevier.com/locate/epsl Chronology of serpentinization: Paleomagnetic evidence for 20 Ma serpentinization of the Oeyama ophiolite, Wakasa, Southwest Japan Yo-ichiro Otofuji a,*, Makoto Fujihara b, Koji Uno InstitfHitryalgr9usaiw bHiruzen Institute for Geology and Chronology, Co., Ltd. 2-5 Nakashima,Naka-ku, Okayama 7038252,Japan Department of Earth Sciences, Okayama University, 3-1-1 Tsushimanaka,Kita-ku, Okayama 700-8530,Japan ARTICLE INFＯ ABSTRACT Article history: Serpentinization of mantle peridotite generates magnetite that is able to acquire a remanent magnetiza- Received 6 June 2020 tion parallel to the ambient geomagnetic field. The paleomagnetic direction of serpentine provides Received in revised form 20 October 2020 insights as to the timing of serpentinization. This study examines paleomagnetic properties of serpentine Accepted 29 October 2020 Available online 12 November 2020 magnetizations with clockwise deflections from the south were observed in the serpentine in the Editor: A. Yin high temperature component (HTC) magnetization with an unblocking temperature (Tub) of ~580°C. Keywords: The mean reversed polarity direction was D = 226.0°, I= -37.8, α95 = 5.9°, and n = 30. The paleomagnetism remaining samples with an HTC of Tub between 400-450°C show the present-day geomagnetic field rock magnetism direction. Rock magnetism reveals that the reversed direction is carried by single-domain (SD) grains serpentine of magnetite, whereas multi-domain (MD) magnetite grains carry the present-day geomagnetic field serpentinization direction. Magnetite grains originated from olivine during serpentinization grew to SD size and acquired ophiolite the chemical remanent magnetizations during a reversed polarity interval. As serpentinization progresses, SD magnetite grains increased to MD magnetite grains and/or were subjected to maghemitization. The mean reversed direction yielded a paleomagnetic pole at 47.4°N, 230.3°E (A95 = 5.7°), which overlaps with the 20 Ma segment of the apparent polar wander path for SW Japan. We conclude that serpentinization occurred in the Oeyama ophiolite at approximately 20 Ma during the clockwise rotation of SW Japan. 2020 Elsevier B.V. All rights reserved. 1. Introduction yielded a U-Pb age of 291 Ma for the Changawuzi ophiolite in southwestern Tianshan (Li et al., 2010) and 400 Ma for the Sude- tic ophiolite in Bohemia (Dubinska et al., 2004): researchers had serpentinized rocks. Serpentinization of mantle peridotite plays a assigned these zircon ages to the timing of serpentinization. How- crucial role in numerous natural processes at mid-ocean ridges, n n n no os se in oceanic subduction zones and ophiolites on land; as such under- remain unclear during serpentinization, the geochronological inter- standing serpentinization is important. Despite its geologic and pretation of the thermal zircon may not always be definitive (Li et tectonic importance, serpentinization dating is extremely difficult al., 2010). As such, many studies aim to determine reliable ages for as minerals suitable for direct isotopic dating are not formed dur- serpentinization. Paleomagnetism is an alternative method to estimate the age ing serpentinization. The rodingitization process occurred contemporaneously with of serpentinization. Magnetite is a significant by-product of the serpentinization (Coulton et al., 1995). The serpentinization age serpentinization of mantle peridotite (e.g., O'Hanley, 1996). The was evaluated by dating rodingitized rocks, which commonly co- transformation of olivine during serpentinization yields magnetite exist with serpentinites (Coleman, 1967). Hydrothermal zircons and serpentine. As a magnetite grain grows through the stable precipitated from the hydrothermal fluid during rodingitization and were associated with serpentinization. Hydrothermal zircons quires a chemical remanent magnetization (CRM) as a natural re- manent magnetization (NRM) parallel to the ambient geomagnetic field. The paleomagnetic pole calculated from the NRM of serpen- * Corresponding author. tine on a tectonic block represents the position at the time of E-mail address: otofuji@kobe-u.ac.jp (Y.-i. Otofuji). serpentinization in its apparent polar wander path (APWP). The https://doi.org/10.1016/jepsl.2020.116656 0012-821X/ 2020 Elsevier B.V. All rights reserved. Y.-i. Otofuji, M.Fujihara and K. Uno Earth and Planetary Science Letters 554(2021)116656 Serpentinites of the Oeyama ophiolite are restricted to one lo- timated at 170 Ma for the Mirdits ophiolite in Albania (Maffione et cality 20 m wide along the Hosomi River in the Wakasa area al., 2013a), despite the sole use of inclination data. (35.39°N, 134.39°E) (Fig. 1). The studied outcrop consists of two n si n lithofacies, units A and B. Unit A is gray to black in color; al- pole position of serpentine with its APWP. As SW Japan has been though serpentinite is altered from dunite, the relict olivine has subjected to clockwise rotation of more than 40° since 100 Ma been preserved, as described by Uemura et al. (1979). In addition e e u a jo do o an pa- to serpentine and olivine, this unit also contains a minor amount leomagnetic poles has been observed in SW Japan between the of clinopyroxene and tremolite. Unit B is gray to green in color, and present-day and 100 Ma (Otofuji and Matsuda, 1987; Uno et al., serpentinite is mainly altered from olivine clinopyroxenite. This 2017). Thus, it is reasonable to determine the timing of serpen- unit mainly consists of serpentine, clinopyroxene, and tremolite. tinization in SW Japan at pre, syn or post-clockwise rotation. In each unit, the serpentinite is not constrained near the fault and The SW Japan arc bears ophiolitic complexes of various ages appears to be a macroscopically homogeneous rock mass. ranging from the early Paleozoic to the Cenozoic (Ishiwatari, 1991, Forty-six block samples were collected from five sites in both 1994; Ishiwatari et al., 2003). The Oeyama ophiolite (Fig. .1) is units. Sampling site ③ was located in unit A, while sites ③, ①, ②, the oldest complex with an age between 566-403 Ma; its main and ? were in unit B. More than eight block samples were col- components are serpentinized peridotite (Kimura and Hayasaka, lected from each site with 1-2 m width. Samples were oriented in 2019). This study focuses on the serpentinite of the Wakasa ul- the field with a tripod-mounted magnetic compass. A declination tramafic complexes (35.39°N, 134.39°E) in the Oeyama ophiolite. correction of D = -9.1° was used, as the sun compass at sites The paleomagnetic pole position of serpentinites at Wakasa was and ② provides a declination of -9.1° ± 0.8°. estimated, and the timing of serpentinization was subsequently ex- amined. 3. Paleomagnetism 2. Geological setting and sampling 3.1. Laboratory procedures SW Japan consists of sequentially developed accretionary com- Cylinders 2.5 cm in diameter were drilled from the 46 block plexes and associated ophiolites, which have developed oceanward u s n p s since the middle Paleozoic (e.g., Moreno et al., 2016; Wallis et al., NRM of the specimens was measured at the laboratory of Okayama University using a spinner magnetometer (Natsuhara Giken SMM- eral elongation of the island arc; the age of ophiolites decreases 85). Thermal demagnetization (ThD) was conducted using a ther- towards the Pacific Ocean. mal demagnetizer composed of a non-inductively wound electric The Oeyama ophiolite, distributed along the north coast of furnace in three cylindrical μ-metal magnetic shields (Natsuhara the SW Japan arc, is the oldest ophiolite in SW Japan (Fig. 1). Giken TDE-91C thermal demagnetizer). The magnetic field mea- It is mainly composed of mainly lherzolite mantle peridotite, sured in the furnace was less than 20 nT. All specimens were clinopyroxene-bearing harzburgite, and dunite, most of which are subjected to progressive ThD in 12-16 heating steps up to a max- altered to serpentine (Ishiwatari and Tsujimori, 2003). Early Pale0- imum temperature of 680°C to detect the presence of hematite. zoic ages have been assigned to the Oeyama ophiolite. The zircon a n s n 206pb/238u ratio for gabbroic rocks yielded ages of 545 and 532 500°℃ (nine steps), and 30°℃ intervals were used from 500 to Ma; these are ascribed to the magmatic crystallization age dur- 680 °C (seven steps). An alternating field demagnetization (AFD) ing oceanic crustal formation (Kimura and Hayasaka, 2019). .The was also applied to the companion specimens for the 17 sam- K-Ar method for hornblende yielded ages between 469 and 403 ples (Table 1). Specimens were subjected to progressive AFD in Ma (Nishimura and Shibata, 1989; Tsujimori et al., 2000). These 10-15 steps at 2.5-20 mT intervals up to 100 mT, with a three- ages are interpreted as the timing of metasomatism during sub- axis tumbler system in three cylindrical μ-metal magnetic shields duction (Kimura and Hayasaka, 2019). The Oeyama ophiolite may (Natsuhara Giken DEM-93). represent a supra-subduction zone ophiolite formed beneath an The demagnetization results were plotted on orthogonal vector intra-oceanic arc (Tsujimori and Itaya, 1999; Ishiwatari and Tsu- diagrams (Zijderveld, 1967) and equal-area projections. Paleomag- jimori, 2003). netic directions were determined by principal component analy- The Wakasa ultramafic complex is one of the exposed areas of sis based on non-anchored fits (Kirschvink, 1980). Site-mean and the Oeyama ophiolite, and the hornblende indicates K-Ar ages be- formation-mean directions were calculated using Fisherian statis- tween 444-469 Ma (Nishimura and Shibata, 1989). It is distributed tics (Fisher, 1953). in small areas along the Hosomi and Kurumino rivers (Uemura et al., 1979) (Fig. 1b). The Wakasa ultramafic complex is in fault con- 3.2. Thermal demagnetization tact with the Renge high-pressure/temperature (P/T) metamorphic belt, and the fault boundary tilts northeastward by 3°-4° (Uemura et al., 1979). The main components of the Wakasa ultramafic com- exception for GY23. For this sample, two specimens were selected; plex are dunite and its derived serpentinite. as such, ThD was conducted on 47 samples. The 47 samples in- The Wakasa ultramafic complex and the Renge high-P/T meta- o s n r morphic belt are overlain by the lower Miocene Tottori Group (Appendix Table). Following the demagnetization of the low tem- and Miocene volcanic rocks (Matsumoto, 1986, 1989; Furuyama et perature component (LTC) at 250°C, a high temperature com- al., 1997). The tilt of the Miocene strata provides paleohorizon- ponent (HTC) with a maximum unblocking temperature (Tub） of tal information for the Wakasa ultramafic complex later than 16 580°C appeared in 37 samples (Fig. 2). A single magnetic com- Ma. The mean tilt from nine localities in the Tottori Group was ponent with a maximum Tub of 400-450°C was observed in the determined to be N37°W/10°E (Ishikawa et al., 2017), and the remaining 10 samples. Miocene volcanic rocks provide N80°E/8°N in attitude (Uemura Three kinds of NRMs were identified for the HTCs; SW, NE, and et al., 1979). Whole rock K-Ar ages of 2.60 ±0.08 Ma and 3.9 ± N remanences (Fig. 3, Table 1, and Appendix Table). Three direc- 1.5 Ma were assigned to the Miocene volcanic rocks (Uto et al., tions and ThD behavior were used to classify 47 samples into five 1994). categories; SWsw, SWne, SWex, NE, and N (Table 1). Y.-i. Otofuji, M. Fujihara and K. Uno EarthandPlanetaryScienceLetters554(2021)116656 (a) Tectonic map 50km Oeyama Ophiolite (b) Wakasa area 1000m Oeyama Ophiolite Plio-Pleistocenevolcanicrocks Miocene volcanic rocks Lower Miocene Tottori Group Fault (c) Sampling locality Fig.1. Simplified tectonic and geological maps together with sketch map of sampling sites. (a) Tectonic framework of the study area showing the Oeyama Belt, and other Paleozoic and Mesozoic geotectonic units (Hida, Suo, Akiyoshi, Maizuru, Ultra-Tamba (UT), Tamba units) (Moreno et al, 2016); (b) simplified geological map of the Wakasa area (Uemura et al., 1979). Renge: Paleozoic geotectonic unit. Whole rock K-Ar ages of 2.60 ± 0.08 Ma and 3.9 ± 1.5 Ma were assigned to the Miocene volcanic rocks (Uto units A and B. (For interpretation of the colors in the figure(s), the reader is referred to the web version of this article.) Y.-i. Otofuji, M.Fujihara and K. Uno Earth and Planetary Science Letters 554(2021)116656 SW Remanence GY21(SWNE) GY74(SWNE) GY71 (SWsw) GY75(SWsw) GY85(SWsw) Dow NE Remanence N Remanence GY36 (SWex) GY53(SWex) GY78 (NE) GY58 (N) Jo =1.7 A/m W_ Up Jo =15.1 A/m E Down EDow Down Fig. 2. Zijderveld plots and normalized demagnetization curves during stepwise thermal demagnetization for samples with SW, NE, and N remanences. Categories of thermal demagnetization behavior are attached to sample numbers. Open (solid) symbols in the Zijderveld plot indicate projection onto a vertical (horizontal) plane. All directions are plotted in geographic coordinates. SW remanence (34 samples) categories. Companion specimens within one block sample (GY23) had different categories; GY23-1 belonged to SWNe whilst GY23-2 Reversed polarity directions with southwesterly declinations was a part of NE category. and upward inclinations were observed in 34 samples. Declinations varied between 170° and 259°, while inclinations varied between 3.3. NRM directions of serpentine in the Oeyama ophiolite -7° and -50° (Fig. 3). Stepwise ThD demonstrated that NRM intensity curves for all Mean directions were calculated for LTCs and HTCs: the results 34 samples showed a sharp intensity decrease at 500-530°C are shown in Fig. 3 and Table 2. NRM directions with a maximum (Fig. 2). A plateau or a gentle peak in the NRM intensity curves angular deviation (MAD) <10° were selected for the mean calcu- occurs up to 500-530°C for 27 of the 34 samples. Only the SW lation. remanence direction (category SWsw) was observed in 19 of these 27 samples throughout the stepwise ThD, whereas a northeasterly (1) Low temperature component direction opposite to the SW remanence direction appeared above 530°C in eight samples (category SWNE). The remaining seven A LTC was observed in 37 samples between room tempera- samples (category SWex) showed swift and exponential decay in ture and 250°C prior to the appearance of the HTC (Table 2). The NRM intensity between room temperature and 250-350°C before mean direction of the LTC was D = 353.5°, I = 47.0° (k = 29.2, the SW remanence direction appeared. Q95 = 4.4°, n = 37), this is statistically indistinguishable from the direction of the axial dipole field (D = 0°,I = 52.8°） and the NE remanence (3 samples) present-day geomagnetic field (International Geomagnetic Refer- Normal polarity directions with northeasterly declinations and ence Field (IGRF): D = 352.3°, I = 50.1°; Thébault et al., 2015). downward inclinations appeared in three samples (category NE); that is, opposite to the SW remanence direction (Fig. 3). The (2) High temperature component stepwise ThD showed an abrupt decrease in NRM intensity at 500-530°C as observed in SW remanence (Fig. 2). NE remanence The mean directions of the HTCs were determined based on the directions were observed in the SWNe category samples, although data of each category, and from data of integrated categories. As these directions were dispersed (Fig. 3). SW remanence was observed in SWsw, SWne, and SWex categories, the mean direction for SW remanence in these three categories Was D = 226.0°, I= -37.8°, k = 21.0, α95 = 5.9°, n = 30. The tilt- N remanence (10 samples): corrected directions are listed in Table 2 using the tilt of the early A northerly direction with a downward inclination was ob- Miocene Tottori Group (N37oW/10E°） or Miocene volcanic rocks ) (S80°W/8°N). They also reveal a clockwise rotated declination. unblocked at 450°C, and their NRM intensity curves were char- The mean direction of the NE remanence was D = 59.0°, I = acterized by exponential decay during stepwise ThD (Fig. 2). 39.4°, k = 72.3, α95 = 14.6°, n = 3, which is antipodal to the SW The remanence measurements indicated that the NRMs of ser- remanence. The mean directions for the NE and SW remanences pentines were not homogeneous: they were variable within a vol- pass the reversal test (McFadden and McElhinny, 1990), with a C ume of ~10o0 cm?. Each sampling site was not consisting of sam- classification at the 95% confidence level (%o = 10.3° < Yc = 18.7°). ples with only one category (Table 1). Samples from sites ? and ? By converting the NE remanence direction to the reversed direc- had three categories, while those from sites ②, ③ and ? had four tion, we calculated a mean direction of D = 227.2°, I= -38.0°, Y.-i. Otofuji, M. Fujihara and K. Uno Earth and Planetary Science Letters 554(2021)116656 Table 1 (a) N (b) N (c) N Sample categories and their rock magnetic characteristics. (a) Category of samples. The sample exhibits three directions: SW, NE, and N remanences. Samples with SW SWsw SWNE* remanence were further classified into three categories (SWsw, SWNE and SWex), (SW remane nce) determined from thermal demagnetization behavior. Samples with NE and N rema- nences were classified as categories NE and N, respectively. Circled numbers and numbers are site and sample names, respectively. Samples with an asterisk were SWNE * (NEremanence) Remanence Category Site (GY) Samples (GY) (a) Categories of samples (d) N (e) N (f) N SW (34) SWsw (19) 33,34 SWex* ⑧ 1118, 83, 85*, 86, 87 ? 71*，75* ② 22,25,26, 28* 15, 51, 52, 54, 56, SWNE (8) 32 NE N 81 73*，74* ② 12, 21*, 23-1, 27 (g) N (h) SWex (7) ? 35, 36*, 82,84*,88* 16, 53* NE(3) NE(3) ③ ⑧ 77*，78* 23-2 SWsw + SWne + SWex SWsw + SWnE + SWex+ NE (SWremanence) (SW remanence) N(10) N(10) ③ 13, 31 ⑧ (i) () (k) N N N 17,72*,76* ② 14, 24 55*, 57, 58* ? Remanence Category Minerals Domain state (b) Rock magnetic characteristics of category SW SWsw Magnetite SD + MD SWNE Magnetite SD AFD AFD AFD + Maghemite MD SWex Magnetite MD + SD SW remanence NE remanence N remanence NE NE Maghemite MD + SD N N Magnetite MD + SD Fig. 3. Equal-area projection of paleomagnetic directions. (a)-(h) Directions after thermal demagnetization for samples of each category and integrated categories. Category: (a) SWsw; (b) and (c) SWNE (SW and NE remanences, respectively); k = 22.0, Q95 = 5.4°, n = 33. The NRM direction of the HTC in this (d) SWex; (e) NE; and (f) N; (g) Integrated categories of SWsw, SWNE and SWex; locality is characterized by a clockwise rotated direction (~50°) (h) categories SWsw, SWNE, SWex and NE; (g) and (h); Black circle: data with MAD <10°; Green circle: data with MAD > 10°; (i)~(k) Directions after alternating field from the south. demagnetization. Remanences: (i) SW; (j) NE; (k) N. Mean directions are calculated The N remanences observed between room temperature and from data with MDF > 10° and are denoted by red stars along with ovals indicating 450 °C in the ThD indicate an in situ mean direction of D = 350.1°, the 95% confidence level (see Table 2). I = 43.1°, k = 29.6, α95 = 9.0°, n = 10. This is statistically indistin- guishable from the axial dipole field and the present-day geomag- Demagnetization curves for category N revealed a distinct ex- netic field. ponential decay, showing low coercivity (Fig. 4). The NRM direc- tions were parallel to the present axial dipole field during AFD 3.4. Alternating field demagnetization go ) P a direction towards the SW was observed above 20 mT in a sam- Following the demagnetization of the low coercivity compo- ple (GY55), suggesting that the SW remanence resides in higher nent, three distinct directions with SW, NE, and N remanences coercivity. were isolated similar to those identified in the ThD (Fig. 4 and Category SWex samples exhibited behavior similar to those of Table 2). The mean directions for the three remanences were sta- s e n ( e s s tistically the same as those isolated during ThD. showed swift and exponential decay curves in intensity, and their The reversed polarity directions of the SW remanence were ob- NRM directions were resultant vectors between the SW and N re- served in SWsw and SWNe categories. A distinct peak between 5 manences. One sample (GY36) exhibited a SW remanence direction and 20 mT in the NRM intensity curve was observed for these and a distinct peak in intensity at 10 mT, after the NRM intensity samples, followed by a concave upward decay curve (Fig. . 4). swiftly decayed. The SW remanence demonstrated high stability in response to The demagnetization curves for category NE samples (GY76-1, the AFD, and more than 15% of the NRM intensity remained at an oe m ( 60 mT. structive field (MDF) was less than 8 mT. The NE remanence was Y.-i. Otofuji, M. Fujihara and K. Uno Earth and Planetary Science Letters 554(2021)116656 Group. Mean direction based on data on samples from each category and from combined categories. n/N: number of samples with MAD < 10° /number of all samples. (inverted): normal direction is inverted to reversed direction. D and I refer to declination and inclination, respectively. k is the Fisherian precision parameter and αgs is the radius Category Polarity Remanence n/N k VGP (°) (°) (N/R) Latitude Longitude (1) ThD results SW SWsw 19/19 223.4 -38.1 6.6 26.7 (0/19) SW SWNE 8/8 226.6 -39.0 14.6 15.3 (0/8) NE SWNE 6/8 37.3 49.2 28.6 6.4 (6/0) SW SWex 3/7 240.4 -31.6 36.6 12.4 (0/3) NE NE 3/3 59.1 39.4 14.6 72.3 (3/0) SW SWsw + SWNE 27/27 224.3 5.9 22.9 (0/27) SW SWsw + SWNE + SWex 30/34 226.0 -37.8 5.9 21.0 (0/30) 230.3°E A95 = 5.70 (Tilt corrected) N37°W/10E° 226.8 -28.2 43.6°N A95 = 5.70 (Tilt corrected) N80°E/8°N 221.4 -33.1 Review A95 = 5.70 SW SWsw + SWNE + SWex + NE(inverted) 33/37 227.2 -38.0 5.5 22.0 (3/30) N N 10/10 350.1 43.1 9.0 29.6 (10/0) IGRF 352.3 50.1 Remanence n/N D 1 A95 Polarity MDF (°) (°) (°) (N/R) (mT) (2) AFD results SW 8/8 233.8 -39.5 6.6 70.8 (0/8) 45.4 ± 10.6 NE 3/3 49.1 50.2 11.6 114.9 (3/0) 6.8 ± 2.7 N 3/3 358.0 40.8 19.0 43.2 (3/0) 4.9 ± 1.4 SW Remanence GY28 (SWsw) Go=7.4(SW/N GY21 (SWnE) EDow NE Remanence N Remanence GY36(SWex) GY53(SWex) GY78 (NE) GY58 (N) GY72 (N) Down Down Dow Fig. 4. Zijderveld plots and normalized demagnetization curves during stepwise alternating field demagnetization for samples with the SW, NE and N remanences defined by onto a vertical (horizontal) plane. All directions are plotted in geographic coordinates. definitely observed between 0 and 15 mT. However, following the S = (3.3 - d)/0.785. The bulk density of a sample was obtained demagnetization of NE remanence, the SW remanence emerged in using the buoyancy method from mass and volume measurements the high-field AFD. The NRM direction of GY76-1 and GY78 moves were made using an electronic force balance (±10-3 g). The de- towards the SW above 15 mT. The NE remanences (D ～ 40°, I ~ d pstl t as na n nds jo an 50°) were also observed in the category SWNe samples (GY21 and Table. GY74) between 0 and 15 mT before the SW remanence was iso- Sample densities in the sampling locality ranged between lated above 20 mT. 2.41-3.21 g/cm?, corresponding to a range between 99% and 12% in serpentinization degree values. Sample GY72 demonstrated an 4. Degree of serpentinization and microscopic observation extraordinarily large serpentinization degree of 114%. Serpentiniza- tion was heterogeneous at each site and sample. The serpentiniza- The degree of serpentinization (S) was calculated from bulk tion degree varied from 19% to 114% within sampling site ? and density (d) using the equation formulated by Miller et al. (1997), from 30% to 59% for sample GY87. Y.-i. Otofuji, M. Fujihara and K. Uno EarthandPlanetaryScienceLetters554(2021)116656 Category SWsw SWNE uen Frequ 20 40 60 80 100 10 SWex requend 工 20 40 60 80 100 Frequency NE 10 10 60 80 100 1 Frequeng 20 40 60 80 100 Fig. 6. Optical photo micrographs of open images. (a) Image for GY71 with a rel- S(%) atively low serpentinization degree (56%). Serpentine and tremolite occurred be- tween the crystal grains of coarse and densely packed pyroxene; (b) image for GY73 with a relatively high serpentinization degree (75%). Coarse pyroxene crystals sur- Fig.5. Frequency distribution of the degree of serpentinization (S). Frequency distri- ao'xodo x dias pe an po se au q pn bution is depicted for five categories. opaque (magnetite), Srp: serpentine, Tr: tremolite. Frequency distributions of the degree of serpentinization are 5. Rock magnetism shown for each category in Fig. 5. Category N samples showed the strongest degrees of serpentinization of >75%, and a degree >70% 5.1. Thermomagnetic analysis is indicated for category SWex samples with the exception of one data point (GY82; 51%). Category NE samples showed a weak de- gree of 31-39%. Although the degree was distributed over a wide Thermomagnetic analyses (high-field magnetization as a func- range between 12% and 99% for 27 samples with the SWsw and tion of temperature) were conducted on one to three samples SWNE categories, two-thirds of samples (21 samples) showed lower across five categories using a laboratory-made magnetic balance degrees of <66%. at Kyoto University (Fig. 7). Samples were heated and cooled at Difference in microstructures of serpentinite was recognized by 8°C/min in air with a direct current (DC) field between 110-250 the degree of serpentinization (Fig. 6). Microstructures were ob- mT. The Curie temperatures (Te) were determined using the inter- served under transmitted and reflected light for eight serpentinite secting tangents method (Prévot et al., 1983). samples from site ? of unit B in the outcrop. The textures of sam- All samples had a Tc between 553 and 582°C, suggesting the ples with fairly low serpentinization degrees (GY71, GY74, GY75, presence of magnetite or partial to completely maghemitized mag- ( netite. Magnetite was expected from a smooth decrease in magne- of serpentine between coarse, densely packed clinopyroxene crys- tization with increasing temperature up to Tc, with the exception tals. Coarse magnetite grains (100-500 μm) was observed between of three samples (GY24, GY36, and GY77). grains of clinopyroxene and olivine, while fine magnetite grains Category NE sample GY77 exhibited an inflection point at ap- were present along the cleavage in clinopyroxene and olivine. On proximately 350°C, while an inflection point at approximately the other hand, the textures of samples with a greater degree of 500°C was observed in samples GY36 (category SWex) and GY24 serpentinization (GY72, GY73, and GY76; 75 ～ 114%) were char- (category N). A small tail above 580 °C was observed in GY78 (cat- egory NE) and GY21 (category SWne). These features are possibly gregates of serpentinite. Coarse magnetite grains up to 500 μm due to the presence of maghemitized magnetite. Was sparsely distributed, and fine magnetite grains occurred along Lower magnetization during cooling than heating was observed cracks or formed veins. There was no maghemite identified in any for all samples (Fig. 7), indicating that oxidation and/or breakdown of the eight samples. of the original magnetic grains occurred during heating in air. Y.-i. Otofuji, M. Fujihara and K Uno Earth and Planetary Science Letters 554 (2021) 116656 (a)SWsw GY71 GY75-1 0.8 0.8 0.6 0.6 0.4 0.4 02 0.2 200 400 600°℃ 200 400 600 °C (b)SWE GY21 GY73 0.8 0.6 0.6 0.4 0.4 0.2 0.2 200 400 600℃ 200 400 600°C (c)SW GY36 0.8 0.6 0.4 0.2 200 400 600°℃ (d) NE GY77 GY78 0.8 0.6 0.6 0.4 0.4 0.2 0.2 200 400 O009 200 00 O009 (e) N GY72 GY76 0.8 0.6 9°0 0.4 0.4 0.2 0.2 200 400 600°C 200 400 600°℃ Fig. 7. Strong field thermomagnetic analysis for samples of each category in air condition. Red and blue lines indicate the heating and cooling curves, respectively. 5.2.Mediandestructivefieldofalternatingfielddemagnetization remanences) appeared during the AFD, the MDF was examined for each remanence (Table 2). AFD provides the MDF, which supplies information on the mag- The MDF for SW remanence was measured from SWsw and netic coercivity of samples.As three remanences (SW, NE,and N SWNE category samples.Their values were larger than 30 mT and Y.-i. Otofuji, M. Fujihara and K. Uno Earth and Planetary Science Letters 554(2021)116656 (a) sW remanence of coarse-grained MD with respect to SD grains (Harrison et al., GY21 (SWnE) GY75 (SWsw) 2018). S-17 % The FORC diagram of a sample (GY72; category N) with N rema- 40 40 nence is indicative of a feature for predominantly coarse-grained 20 MD (Pike et al., 2001). The diagram shows a considerable vertical 20 spread (along the Hu axis). There is also a subtle central-ridge-like (1W)H (1w) 0 0 distribution along the Hc axis (Hu = 0), extending to ～ 40 mT. The FORC distribution function abruptly decreased with increasing Hc. -20 -20 5.4. Day plot -40 -40 The hysteresis parameters were also measured using the same 0 20 40 60 80 100 20 40 60 80 100 1 instrument used for FORC. The hysteresis parameters for eight H(mT) H(mT) samples were represented in a Day plot (Fig. 9), with saturation (b) NE remanence (c) N remanence remanent magnetization/saturation remanence (Mrs/Ms) versus re- GY72 (N) manent coercive force/coercive force (Hcr/Hc), to obtain informa- GY78 (NE) tion on the magnetic domain state and magnetic grain size. 40 S=31 % 40 S=114 % The Mrs/Ms measured values for all samples ranged from 0.05 00 to 0.16, while the ratio of Her/Hc ranged from 1.6 to 3.7. In terms 20 20 of hysteresis ratios of Mrs/Ms versus Hcr/Hc, almost all samples (1W)H E fell in the pseudo-single-domain (PSD) region (Day et al., 1977). 0 0 A variable proportion of SD and MD grains was anticipated, as 工 -20 -20 data are distributed along the theoretical SD + MD mixing curve -40 -40 a larger SD fraction with respect the MD fraction than the N rema- nence samples. 0 20 40 60 80 100 60 80 H(mT) H(mT) 5.5.ModifiedLowrie-Fuller test Fig. 8. First-order reversal curves (FORC) diagrams (SF = 3) for samples with SW, The modified Lowrie-Fuller test (Lowrie and Fuller, 1971) was NE, and N remanences (Harrison and Feinberg, 2008). VARIFORC was not used and applied to three samples with a SW or a N remanence, respec- first-point was removed during creating of the FORC diagrams. SW remanence; GY21 of category SWNE and GY75 of category SWsw, NE remanence; GY78 of cate- tively (Fig. 9). An anhysteretic remanent magnetization (ARM) was gory NE, and N remanence; GY72 of category N. S is the degree of serpentinization. oi e a i e while IRM was produced in a 100 mT field. SW remanence; GY71, GY75, GY85 (category SWsw) the MDF of GY75 was as large as 59 mT: The mean MDF is 45.4 ± The degrees of serpentinization for three samples vary between 10.6 mT. The N remanence in category N samples had the low- 19 and 56%. est MDFs, less than 6.5 mT: The mean MDF is 4.9 ± 1.4 mT. The ARM demagnetization curve was considerably more resis- As magnetite is a magnetic carrier for samples with SW and N tant than the IRM curve throughout AFD. The presence of a signif- remanences deduced from thermo-magnetic analysis, the MDF is icant amount of SD-sized grains was demonstrated by the Lowrie- an inverse proxy for magnetic grain size. SD grains (30 mT < Fuller test (Bailey and Dunlop, 1983; Halgedahl, 1998). Firstly, the MDF) were predominant in samples with the SW remanence, while MDF of ARM was more than twice that of IRM. Secondly, the ARM MD grains (MDF < 10 mT) dominated samples with N rema- decay curve was initially convex upward and then convex down- nence. ward, which is uncharacteristic of large MD behavior (Dunlop and The MDFs of the NE remanence were less than 8.5 mT, while Ozdemir, 1997). However, the presence of a small amount of MD the mean MDF was 6.8 ± 2.7 mT. MD-like AFD curves were also grains was also expected from the exponential decay curves for observed. IRM. This AFD behavior has been interpreted as being “bimodal' with SD and MD grains in the grain size distribution, as postulated 5.3. First-Oder Reversal Curves (FORCs) by Dunlop (1983). N remanence; GY58, 72, 76 (category N) The degrees of serpentinization for three samples vary between First-order reversal curve (FORC) experiments for seven sam- 75 and 114%. ples were carried out using the Princeton Measurements Micro- The IRM demagnetization curve for sample GY76 was similar Mag model 3900 vibrating sample magnetometer (VSM) of the to the ARM curve in shape and the MDF of the IRM was larger Paleo and Rock magnetism Laboratory at the Kochi Core Center, than the MDF of the ARM, indicating the presence of large MD-size Kochi University. During FORC measurements, the number of FROC grains. Samples GY58 and GY72 showed exponential-like demagne- m nd d a n i ds s tization curves for the ARM accompanied by a low-coercivity IRM 1.3 mT and 100 ms, respectively. FORC diagrams were created us- with exponential decay. The AFD behavior for these samples has ing FORCinel (Harrison and Feinberg, 2008) (Fig. 8). been interpreted as being “bimodal' in the grain size distribution The FORC distribution of sample GY21 (category SWNe) with a (Dunlop, 1983), that is, SD and MD grains. The MDF of ARM does SW remanence indicates dominant fine-grained SD magnetite. The not exceed twice that of IRM, suggesting that the MD fraction is presence of well-dispersed SD grains was confirmed by a strong larger than of that a sample with the SW remanence. "central ridge" distribution along the Hc axis up to ~100 mT with no vertical spread (Roberts et al., 2014; Harrison et al., 2018). 6. Discussion In samples GY75 (category SWsw） and GY78 (category NE), the central ridge along the Hc axis had decrease and vertical spread The rock magnetism and paleomagnetism provide informa- along the Hu axis had risen, suggesting increase in the proportion tion on magnetic carriers and magnetic grain size. In this study, 9 Y.-i.Otofuji,M.Fujihara andK.Uno EarthandPlanetaryScienceLetters554(2021)116656 (a) Day plot 0.6 SD 0.5 0.4 ● SW remanence ●NE remanence 0.3 ●N remanence M PSD 0.2 0.1 7872 MD- 2 3 5 Hcr/Hc (b)Lowrie-FullerTest SW remanence N remanence GY71(SWsw) GY58(N) 0.6 0.6 0.4 0.2 20 40 60 80 20 40 60 80 alternating field (mT) alternating field (mT) GY75(SWsw) GY72(N) 0.6 0.4 0.2 0.2 20 40 60 80 20 40 60 80 alternating field (mT) alternating field (mT) GY85(SWsw) GY76(N) 0.8 0.6 norr 0.4 20 40 60 80 20 40 60 80 alternating field (mT) alternating field (mT) Fig. 9. Rock magnetic results. (a) Hysteresis data for samples with the SW, NE, and N remanences. Mrs, saturation remanent magnetization; Ms, saturation magnetization; Hcr, remanent coercive force; Hc, coercive force. Single-domain (SD), pseudo-single-domain (PSD) and multi-domain (MD) fields after Day et al. (1977). Almost all samples fall in the PSD region. The dotted curve represents a SD-MD theoretical mixing curve for magnetite (Dunlop, 2002); (b) modified Lowrie-Fuller test: SW remanence (category SWsw GY71, 75, 85) versus N remanence (category N; GY58, 72, 76). The IRM (black line with circles) was produced by applying a 100 mT DC feld of and the ARM (red line with triangles) was produced by applying a 100 μuT DC biasing field in a peak 100 mT AC field. 10 Y.-i. Otofuji, M. Fujihara and K. Uno Earth andPlanetary ScienceLetters554(2021)116656 we propose an evolutionary aspect of magnetic carriers in the tion of magnetite (Andreani et al., 2013). Serpentinization may be Oeyama ophiolite, and then estimate the timing of serpentiniza- described in a simple form as (e.g., Evans, 2008; Klein et al., 2013): tion based on the paleomagnetic direction of the magnetic car- rier, which records the Earth's magnetic field during serpentiniza- MgFe olivine + H20 tion. → MgFe serpentine + magnetite + MgFe brucite + H2 (1) 6.1. Magnetic carriers of remanent magnetizations Magnetite is also formed from secondary minerals, such as, brucite and Fe-bearing serpentine (Bach et al., 2006; Frost et al., 2013; The NRM of the serpentine from the Oeyama ophiolite was Klein et al., 2014; Schwarzenbach et al., 2016) as follows: carried by magnetite or maghemitized magnetite. The MDF was less than 59 mT for samples of all remanences, suggesting that FeMg brucite + SiO2 their magnetic carriers have a coercivity much lower than that of hematite. Almost all samples had Tc between 553 and 582°C → FeMg serpentine + magnetite + H2O + H2 (2) (Fig. 7). These results indicate that magnetite is the predominant FeMg brucite → magnetite + H2O + H2 (3) magnetic carrier. Possible magnetic carriers in samples with the five categories are listed in Table 1. Fe serpentine → magnetite + SiO2 + H2O + H2. (4) SW remanence in category SWsw and SWne samples is carried s e n si s Although the specific reactions are complex, serpentinization pro- Js-T curve during heating (Fig. 7). The MDF analysis, FORC di- duces magnetite. agrams and Lowrie-Fuller test (Figs. 8 and 9) indicate that the The Fe released from olivine, brucite and serpentine in ordi- magnetization of category SWsw and SWNE samples was carried nary serpentinized peridotites is precipitated in situ as fine dis- by SD magnetite grains. Therefore, the SW remanence originated crete magnetite grains. As magnetite crystals grow and the crystal from magnetite SD grains. size of superparamagnetic magnetite increases to a stable SD size The MDF and FORC diagram (Fig. 8) indicate that magnetization in a magnetic field, the SD magnetite grains acquire grain-growth of the category N samples was carried by MD magnetite grains; the Crystallization remanent magnetization (CRMmagnetite). results from the Lowrie-Fuller test support this conclusion (Fig. 9). Two major alteration processes are anticipated for the evolution The presence of MD grains was also inferred from the exponen- of SD magnetite during serpentinization in a temperature envi- tially swift decay of the NRM intensity during the stepwise ThD ronment lower than ~300°C, that is, growing into MD grains or s ) maghemitization. First, as serpentinization progresses, the mag- (Dunlop and Ozdemir, 200o). The N remanence was ascribed to netite crystals grow and the magnetite grain size increases from the magnetization of MD magnetite grains. the SD to the MD stage in grain size (Maffione et al., 2013a). The Magnetization for category SWex samples is likely to consist MDF analysis demonstrates that samples with the N remanence - n m s an pe n s m s as jo are magnetically softer than those with SW remanence. Based on nence. NRM behavior during the ThD was an intermediate between FORC diagrams and the Lowrie-Fuller test (Figs. 8 and 9b), the the SW and N remanences; NRM intensity shows a swift decay at proportion of the MD/SD magnetite grain size ratio in the N re- 500-530°C following exponential decay by 100-250 °C. n Maghemitized magnetite is the most plausible candidate for NE magnetite grains are easily able to acquire the present-day mag- remanence. The presence of maghemitized magnetite was iden- netic field as a VRM (Shimizu, 1960; Dunlop and Ozdemir, 1997), tified in a category NE sample (GY77). Although the Tc of this the NRM of samples with N remanence are parallel to the present- sample is 579°C, its Js-T curve shows an inflection point at ap- day geomagnetic field direction. Crystal growth from the SD size proximately 350°C (Fig. 7). This finding is consistent with the magnetite to the MD size magnetite is one form of the evolution inversion temperatures for presumably fine-grained maghemitized of magnetic minerals in serpentine. This estimate is plausible as magnetite (e.g., de Boer and Dekkers, 1996; Dunlop and Ozdemir, category N samples were limited to serpentine with a high degree 1997; Liu et al., 1999), or transition temperatures from less stable of serpentinization (Fig. 5). maghemite into more stable maghemite (Gendler et al., 2005). The Second, an alternative form of evolution is low-temperature thermomagnetic curve also exhibits a small tail just immediately Oxidation (maghemitization), that is, Fe3O4 → yFe2O3. Maghemi- above 580°C in GY78, indicating the existence of maghemitized tization takes place in an environment <200 °C, occurring mainly magnetite (Helgason et al., 1992). A small tail above 580°C was at the crystal surface (Dunlop and Ozdemir, 1997). It should be observed in the category SWNe (GY21) sample. During ThD and noted that magnetite associated with serpentinization forms be- AFD for this sample (Figs. 2 and 4), the NE remanence had a lower tween 150 and 350°C and its effective temperature is approxi- coercivity and a Tub higher than those of magnetite; that is further mately 300 °C (Klein et al., 2009). Therefore, after magnetite forms evidence of maghemitized magnetite carrying the NE remanence during serpentinization, maghemitization may occur in an environ- (Johnson et al., 1975; Goss, 1988; Dunlop and Ozdemir, 1997). The ment with decreasing temperature. MDF 6.8 mT would be ascribed to the SD grains of maghemitized magnetite. 6.3. Origin of NE remanence direction with normal polarity 6.2. Evolution aspect of magnetic carriers of remanent magnetization Maghemitized magnetite in NE and SWNE categories acquired the remanent magnetization of another type of CRM (CRM) asso- The SD grains of magnetite are likely to be the first occur- ciated with maghemitization during the post-serpentinization pe- rence of magnetic grains with stable NRM during serpentinization riod. During serpentinization, the SD magnetite (categories SWsw. as suggested by Bina and Henry (1990) and Maffione et al. (2013a). SWNE and SWex) acquired CRMmagnetite, parallel to the ambient The NRM of the SWsw category is the original remanence during geomagnetic field, and its direction was D ～ 230°, I~ -40° the serpentinization in the Oeyama ophiolite. (Fig. 3). When the originally produced SD magnetite grains altered Serpentinites form via hydration of olivine-rich ultramafic rocks to the maghemitized magnetite during the maghemitization, the between 200-400 °C and 0.3-0.5 kbar accompanied by the forma- maghemitized magnetite grains secondarily acquired the CRM. 11 Y.-i. Otofuji, M. Fujihara and K. Uno Earth andPlanetary ScienceLetters554(2021)116656 O°E (a) (b) NCBpole SWJ12Ma Serpentine pole (rotation corrected) 0°E 70°E SWJ20Ma Rotationpole Sampling area Serpentine pole Sampling area Serpentine pole SWJ 100 Ma SWJ 35-30 Ma SWJ110Ma SWJ90-70Ma SCBpole 180°E 180°E Fig. 10. Stereographic projection of the paleomagnetic pole from the serpentine of the Oeyama ophiolite in the Wakasa area (this study). (a) The serpentine pole position (pink) compared with the APWP of SW Japan (SWJ; 110 Ma to 12 Ma) summarized by Uno et al. (2017); (b) the reconstructed pole position of the serpentine prior to the clockwise rotation of SW Japan (rotation corrected) compared with the Paleozoic pole for NCB (Huang et al., 1999) and SCB (Lin and Fuller, 1990). We propose two potential factors for the acquisition of the 1= -37.8° (αg5 = 5.9°, N = 30), whereas the latter is D = 227.2°, CRM, with a direction opposite to that of the CRMmagnetite; am- I= -38.0° (Q95 = 5.4°, N = 33). Although the two directions bient geomagnetic field or a self-reversal mechanism. The CRMy are statistically indistinguishable, we adopted the SW remanence is acquired during maghemitization in the geomagnetic field direc- directions as the characteristic direction of serpentinization. This tion with normal polarity, which occurred following the reversed characteristic direction yields a pole position of 47.4°N, 230.3°E polarity period. An alternative explanation is the self-reversal phe- with Ag5 = 5.7° for the Oeyama ophiolite. nomenon, in which maghemitized magnetite grows to SD grains We confirm that no serpentinization occurred in the Paleozoic and acquires a CRMy opposite to the direction of the magnetite. when the Oeyama ophiolite occupied the Oeyama area. Because The NE remanence directions of the self-reversal origin may be it is controversial whether the SW Japan arc of the Paleozoic era regarded as equivalent for the SW remanence directions of SD belonged to the NCB or SCB (Isozaki and Maruyama, 1991; Ishi- magnetite in the SWsw, SWNE and SWex categories, although its watari and Tsujimori, 2003), the Oeyama pole at serpentinization polarity is opposite. is compared with those of the NCB and SCB. Paleomagnetic poles Self-reversal phenomena for titanium-poor magnetite are enig- at approximately 450 Ma are constructed from the paleomagnetic matic. The self-reversal of titanomagnetites with relatively high Ti data of NCB and SCB (Fig. 10). The paleopole during the Middle - s o s (io x) Ordovician (472-461 Ma) is estimated at 31.5°N, 327.7°E for the covered by the Ocean Drilling Program (ODP) (Doubrovine and NCB (Huang et al., 1999), while the paleopole during the Silurian Tarduno, 2004, 2006; Cavallo et al., 2010). O'Reilly and Baner- (425 Ma) is at 6.8°N, 195.6°E for the SCB (Lin and Fuller, 1990). jee (1966) theoretically proved that the self-reversal phenomenon As SW Japan experienced the clockwise rotation due to opening of is restricted to oxidized titanomagnetite. On the other hand, the the Japan Sea in the early Miocene (Otofuji et al., 1985; Jolivet et laboratory simulation of maghemitization oxidized from SD-size al., 1994), the geomagnetic pole for the Oeyama ophiolite must be titanium-poor magnetites demonstrates that the maghemite ac- calculated from the paleoposition before the rotation of SW Japan. quires the CRM direction parallel to the original NRM direction SW Japan is reconstructed by counter-clockwise rotation on a Eu- (Johnson and Merrill, 1972; Ozdemir and Dunlop, 1985). ler pole (34°N, 129°E) by 42° (Otofuji and Matsuda, 1987). Then Although we have proposed either mechanism due to the geo- the geomagnetic pole calculated from the reconstructed position magnetic field or self-reversal for these discovered NE remanence of SW Japan shifts to 73.4°N, 302.0°E. This reconstructed pole po- directions with normal polarity, we were unable to determine the sition is far from the Paleozoic paleomagnetic poles for the NCB responsible origin. Further studies are required to determine the and SCB, suggesting that serpentinization of the Oeyama ophiolite magnetic nature of serpentines. is not ascribable to the Paleozoic phenomenon. Serpentinization occurred during the Cenozoic. The geomag- 6.4. When did serpentinization occur? netic pole position for the Oeyama ophiolite is compared with the APWP of SW Japan. Uno et al. (2017) reported the APWP of SW The timing of serpentinization is estimated on the basis of a Japan between 110 and 12 Ma (Fig. 10). The geomagnetic pole comparison between the NRM direction of the Oeyama ophiolite for the Oeyama ophiolite coincides with the 20 Ma pole position and the paleomagnetic directions in SW Japan. This study proposes (51.3° N, 221.3° E with A95 = 6.7°) of SW Japan. two characteristic paleomagnetic directions for serpentinization in The effect of structural correction is examined for the pole po- the Oeyama ophiolite (Fig. 3). First, the mean SW remanence di- sition of the Oeyama ophiolite. Using the tilt of the early Miocene rection is calculated from the data for categories SWsw, SWNE, Tottori Group and Miocene volcanic rocks, the tilt-corrected pale- and SWex of the SD grains in the magnetite. The second direc- omagnetic direction yields a pole position at 43.6° N and 238.1°E tion is based on the SW remanence directions of categories SWsw, or 44.0°N and 222.9°E respectively. The tilt-corrected paleomag- SWNE, SWex and NE, assuming that the NE remanence direction netic poles show few large deviations from the pole at 20 Ma of is a contemporaneous record of the geomagnetic field in the SW ' ms o sod 1 oi-o yde nou op ue ud s remanence direction. The former mean direction is D = 226.0°, Therefore, we conclude that the timing of 20 Ma is a plausible 12 Y.-i. Otofuji, M. Fujihara and K. Uno Earth and Planetary ScienceLetters 554(2021)116656 estimate for serpentinization in the Oeyama ophiolite. The serpen- analysis Koji Uno: Conceptualization, Methodology, Investigation, tinization at the Wakasa area occurred until the opening of the Formal analysis, Writing Japan Sea ceased at 15 Ma. The serpentinization age for the Oeyama ophiolite of 20 Ma is Declaration of competing interest (oo) e na osi q psns Aisnoad u no yn A U-Pb age of 472 ± 8.5 Ma was observed for zircon crystals in a All authors have no conflicts of interest in academic research. jadeite vein in the Osayama serpentine mélange of a part of the Oeyama ophiolite, located to the west of the Wakasa area (Fig. 1). Acknowledgements Tsujimori et al. (2005) recognized that this zircon age in jadeite as- sociated with serpentine is the timing of the interaction between The authors would like to thank Naoto Ishikawa (Kyoto Uni- alkaline fluid and ultramafic rocks; thus, these data are not direct versity) for providing the laboratory facilities. We thank Yuhji Ya- time constraints of serpentinization. We conclude that the Pale- mamoto for FORC and hysteretic parameter measurements at the ozoic peridotite in the Oeyama ophiolite altered to serpentine at Paleo and Rock magnetism Laboratory at the Kochi Core Center, approximately 20 Ma. Alternatively, 20 Ma serpentinization is the Kochi University. We would like to acknowledge the useful sugges- most recent phenomenon in multi occurrences of serpentinization tions from Masako Miki. We thank K.P. Kodama and M.J. Dekkers since the ophiolite formed. Magnetite originated in older phases of for their careful and useful reviews. This research work was partly e sn nn g pe sua a on dn smon uonzd s n ( ) -- prn parallel to the present-day geomagnetic field. istry of Education, Culture, Sports and Technology (MEXT). n on od si e o j d o do a ne s no u s Appendix A. Supplementary material the Japan Sea (Otofuji et al., 1991; Jolivet et al., 1994; Baba et al., 2007). Serpentinization at pre or syn-event of the Japan Sea Supplementary material related to this article can be found on- s o s e ss oe si on line at https: //doi.0rg/10.1016/j.epsl.2020.116656. tinized ultramafic rocks in the Hayachine tectonic belt, northern Kitakami Terrane, northeast Japan (Otofuji et al., 2003). Although References ocean ridges, seafloor detachment faults or subduction settings Andreani, M., Munoz, M., Marcaillou, C., Delacour, A., 2013. μXANES study of (e.g., Coulton et al., 1995; Dubinska et al., 2004; Deschamps et al., iron redox state in serpentinization during oceanic serpentinization. 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Jolivet, L, Tamaki, K, Fournier, M., 1994. Japan Sea, opening history and mechanism: Thébault, E., Finlay, C.C., Beggan, C.D., Alken, P, Aubert, J., Barrois, O., Bertrand, F, a synthesis. J. Geophys. Res. 99, 22,237-22,259. Bondar, T., Boness, A., Brocco, L., Canet, E., Chambodut, A., Chulliat, A., Coisson, Kimura, K, Hayasaka, Y, 2019. Zircon U-Pb age and Nd isotope geochemistry of P., Civet, F, Du, A., Fournier, A, Fratter, I, Gillet, N, Hamilton, B., Hamoudi, M, latest Neoproterozoic to early Paleozoic Oeyama ophiolite: evidence for oldest Hulot, G., Jager, T, Korte, M., Kuang, W., Lalanne, X., Langlais, B., Leger, J.-M. MORB-type oceanic crust in Japanese accretionary system and its tectonic im- Lesur, V., Lowes, FJ, Macmillan, S., Mandea, M., Manoj, C., Maus, S., Olsen, N. plications. Lithos 342-343, 345-360. Petrov, V., Ridley, V., Rother, M, Sabaka, TJ, Saturnino, D., Schachtschneider, R, Kirschvink, J.L., 1980. The least-squares line and plane and the analysis of palaeo- Siro, O., Tangborn, A., Thomson, A., Toffner-Clausen, L., Vigneron, P, Wardinski, magnetic data. Geophys. J Int. 62, 699-718. 1., Zvereva, T, 2015. International geomagnetic reference field: the 12th genera- Klein, FE, Bach, W., Jons, N., McCollom, T., Moskowit, B., Berquo, T, 2009. Iron par- tion. Earth Planets Space 67, 79. https://doi.org/10.1186/s40623-015-0228-9. d hssae po eds a unu e o Tsujimori, T, Itaya, T, 1999. Blueschist-facies metamorphism during Paleozoic from 15°N on Mid-Atlantic Ridge. Geochim. Cosmochim. Acta 73, 6868-6893. orogeny in southwestern Japan: phengite K-Ar ages of blueschist-facies tectonic Klein, F, Bach, W., McCollom, T.M., 2013. Compositional controls on hydrogen gen- blocks in a serpentinite melange beneath early Paleozoic Oeyama ophiolite. Isl. eration during serpentinization of ultramafic rocks. Lithos 178, 5-69. Arc 8, 190-205. Klein, F, Bach, W., Humphris, S.E., Kahl, W.A, Jons, N, Moskowitz, B., Berquo, TS. Tsujimori, T., Nishina, K., Ishiwatari, A., Itaya, T., 2000. 443-403 Ma kyanite-bearing 2014. Magnetite in seaflor serpentinite-some like it hot. Geology 42, 135-138. epidote amphibolite from the Fuko Pass metacumulate in the Oeyama area, the Li, X-P., Zhang, L-F, Wilde, S.A., Song, B., Liu, X-M, 2010. Zircons from rodingite in Inner Zone of southwestern Japan. J. Geol. Soc. Jpn. 106, 646-649. the Western Tianshan serpentinite complex: mineral chemistry and U-Pb ages Tsujimori, T., Liou, J.G., Wooden, J., Miyamoto, T, 2005. U-Pb Dating of large zircons define nature and timing of rodingitization. Lithos 118, 17-34. in low-temperature jadeitite from the Osayama Serpentinite Melange, South- Lin, J.L, Fuller, M, 1990. Palaeomagnetism, North China and South China collision, west Japan: insights into the timing of serpentinization. Int. Geol. Rev. 47, and the Tan-Lu fault. Philos. Trans. R. Soc. Lond. A 331, 589-598. 1048-1057. Liu, X.M., Hesse, P, Rolph, T., 1999. Origin of maghemite in Chinese loess deposits: Uemura, F, Sakamoto, T, Yamada, N., 1979. Geology of the Wakasa District. Quad- aeolian or pedogenic. Phys. Earth Planet. Inter. 112, 191-201. rangle series, scale 1:50,000. Geol. Surv. Jpn., 91 p. 14 Y.-i. Otofuj, M. Fujihara and K Uno Earth and Planetary Science Letters 554 (2021) 116656 Uno, K., Furukawa, K., Hatanaka, Y., 2017. An analysis of apparent polar wander path Wallis, S.R., Yamaoka, K., Mori, H., Ishiwatari, A, Miyazaki, K., Ueda, H., 2020. The for southwest Japan suggests no relative movement with respect to Eurasia dur- basement geology of Japan from A to Z. Isl. Arc 29, e12339. https://doi.org/10. ing Cretaceous. Phys. Earth Planet. Inter. 267, 1930. 1111/iar.12339. Uto, K, Tagami, T, Uchiumi, S., 1994. K-Ar and fission-track dating on volcanic rocks Zijderveld, J.D.A., 1967. A.C. demagnetization of rocks: analysis of results. In: of Pliocene Teragi Group from eastern San'in region, Southwest Japan. J. Geol. Collinson. D.W. Creer. KM. Runcorn, S.K. (Eds.), Method in Palaeomagnetism. Soc. Jpn. 100, 787798. Elsevier. Amsterdam. p. 254286. 15
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Otofuji (2021) chronology of serpentinization oeyama ophiolite.txt
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Seismological features of island arc crust as inferred from
recent seismic expeditions in Japan
Takaya Iwasakia,*, Toshikatsu Yoshiia, Tanio Itob, Hiroshi Satoa, Naoshi Hirataa
aEarthquake Research Institute, University of Tokyo, Yayoi 1-1-1, Bunkyo, Tokyo, 113-0032, Japan
bFaculty of Science, Chiba University, 1-33 Yayoi-cho, Inage-ku, Chiba, 263-8522, Japan
Received 6 January 2001; received in revised form 28 July 2001; accepted 8 August 2001
Abstract
Crustal studies within the Japanese islands have provided important constraints on the physical properties and deformation
styles of the island arc crust. The upper crust in the Japanese islands has a significant heterogeneity characterized by largevelocity variation (5.5–6.1 km/s) and high seismic attenuation (Qp = 100–400 for 5–15 Hz). The lateral velocity change
sometimes occurs at major tectonic lines. In many cases of recent refraction/wide-angle reflection profiles, a ‘‘middle crust’’
with a velocity of 6.2–6.5 km/s is found in a depth range of 5–15 km. Most shallow microearthquakes are concentrated in theupper/middle crust. The velocity in the lower crust is estimated to be 6.6–7.0 km/s. The lower crust often involves a highlyreflective zone with less seismicity, indicating its ductile rheology. The uppermost mantle is characterized by a low Pn velocity
of 7.5–7.9 km/s. Several observations on PmP phase indicate that the Moho is not a sharp boundary with a distinct velocity
contrast, but forms a transition zone from the upper mantle to the lower crust. Recent seismic reflection experiments revealedongoing crustal deformations within the Japanese islands. A clear image of crustal delamination obtained for an arc–arc
collision zone in central Hokkaido provides an important key for the evolution process from island arc to more felsic continental
crust. In northern Honshu, a major fault system with listric geometry, which was formed by Miocene back arc spreading, wassuccessfully mapped down to 12–15 km.
D2002 Elsevier Science B.V. All rights reserved.
Keywords: Seismological features; Island arc crust; Seismic expeditions
1. Introduction
The Japanese islands are composed of several
island arcs developed along subduction zones alongthe eastern margin of the Asian Continent (Fig. 1) .
They are under complex tectonic circumstances domi-nated by plate subduction, accretion, back arc spread-ing and arc–arc collision. Major islands of Hokkaido,Honshu, Shikoku and Kyushu are geologically divided
into two arcs of NE Japan and SW Japan. The NE
Japan Arc, which includes northern Honshu and west-
ern Hokkaido, is overriding the subducting PacificPlate. The SW Japan Arc consists of the western halfof Honshu and Kyushu, beneath which the PhilippineSea plate is subducting. During most of Mesozoic toEarly Miocene time, these arcs were situated along thesubduction zone in the eastern margin of the AsianContinent, and rotated to their present locations by
back arc opening of the sea of Japan 20–14 Ma (e.g.
0040-1951/02/$ - see front matter D2002 Elsevier Science B.V. All rights reserved.
PII: S 0040-1951(02)00134-8*Corresponding author. Fax: +81-3-5689-7234.
E-mail address: iwasaki@eri.u-tokyo.ac.jp (T. Iwasaki).www.elsevier.com/locate/tectoTectonophysics 355 (2002) 53–66
Otifuji et al., 1985 ). The extensional stress regime
responsible for the back arc opening was changed intocompressional stress at 4 Ma in NE Japan and at 2–3Ma in SW Japan (e.g. Sato, 1994 ). Seismic activity in
and around these two arcs is extremely high ascharacterized by M8 class interplate earthquakes andM7 class intraplate earthquakes. Recent microseismic
observations revealed that most of the crustal earth-quakes are concentrated in a depth range of 0–15 km,w h i c hs h o w sam a r k e dc o n t r a s tw i t hav e r yl o wseismic activity in the lower crust (Ito, 1993, 1999) .
Deeper events in NE Japan are forming double seismic
Fig. 1. Tectonic map in and around Japan. Major profile lines are also indicated with shot points (stars). Broken lines show tectonic lines. A –A V:
Oga–Kesennuma profile (Yoshii and Asano, 1972; Okada et al., 1979) .B – BV: Shakotan–Erimo profile (Okada et al., 1973) .C – CV: Atsumi–
Noto profile (Aoki et al., 1972) . a: 1990 profile (Iwasaki et al., 1994) . b: 1997 profile (Iwasaki et al., 2001) .bV: Area of 1998 experiment.
c: 1999–2000 profile (Iwasaki et al., 2000a,b; Moirya et al., 2000) . TTL: Tanagura Tectonic Line. KTL: Kanto Tectonic Line, ISTL: Itoigawa–
Shizuoka Tectonic Line. ATL: Akaishi Tectonic Line.T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 54
planes within the subducting Pacific Plate down to 200
km (e.g. Hasegawa et al., 1978 ). In the SE Japan, such
seismic planes are less developed and limited at adepth shallower than 100 km (e.g. Ishida, 1992 ).
Off the southern part of central Honshu, the Japan
Trench, Nankai Trough and Izu-Bonin Trench meettogether forming a triple junction to produce complex
structures. The Izu-Bonin Arc on the Philippine Sea
plate is colliding to Honshu from the south, causingextremely high volcanic and seismic activities. Thiscollision is also responsible for the formation of asyntaxis by the Itoigawa–Shizuoka Tectonic Line(ISTL) and the Kanto Tectonic Line (KTL), the latterof which is considered to separate SW Japan to NEJapan. There exists another arc–arc collision zone in
central Hokkaido (Fig. 1) . Since Miocene, the Kuril
Forearc, now occupying the southernmost part ofeastern Hokkaido, has been collided against the west-ern Hokkaido (the NE Japan Arc) to form the HidakaCollision Zone.
2. Overviews of previous refraction/wide-angle
reflection experiments
Crustal studies of the Japanese islands using con-
trolled seismic sources started in the 1950s. Expedi-tions in the early stage (1960s) were focused onelucidating large-scale structural variation of the crustand upper mantle across the Japanese islands (Aoki et
al., 1972; Yoshii and Asano, 1972; Okada et al., 1973,1979) . Particularly, the horizontal change in Pn veloc-
ity was intensively investigated. Fig. 2 shows a crustal
model across the NE Japan Arc (Oga-Kesennumaprofile; profile A–A Vin Fig. 1) by Yoshii and Asano
(1972) andOkada et al. (1979) . The most important
feature of this model is an anomalously low Pnvelocity ( f7.5 km/s) beneath the NE Japan Arc.
The Pn velocity becomes high either to the forearc or
back arc side. The crust of the NE Japan Arc is 30 kmthick, composed of upper and lower layers with avelocity of 5.9 and 6.6 km/s, respectively. In thewestern part of this profile, there is a remarkablecrustal thinning probably related to the Mioceneopening of the sea of Japan. Relatively low Pnvelocities of 7.5–7.9 km/s are also reported by crustal
studies of Atsumi-Noto and Shakotan-Erimo profiles
(Aoki et al., 1972; Okada et al., 1973 ;profiles B–B V
and C–C V, respectively, in Fig. 1).
Seismic experiments in the 1980s revealed upper
crustal heterogeneities. The P-wave velocity of theuppermost crust is slightly lower than 6 km/s (5.8–5.9 km) in average, but shows a significant regionalvariation ranging from 5.5 to 6.1 km/s (Yoshii, 1994) .
A drastic velocity change sometimes occurs across
major tectonic lines of ISTL and ATL ( Ikami et al.,
1986; Matsu’ura et al., 1991; Sakai et al., 1996;Takeda, 1997 ; see Fig. 1 ). Seismic refraction studies
in the 1990s on longer profiles (150–200 km) sampledcrustal velocities down to a 15–20 km depth. In theeastern part of NE Japan (profile ‘a’ in Fig. 1) , the
5.85–6.15 to 6.15–6.25 km/s uppermost crust is
Fig. 2. Crustal model across NE Japan (Oga–Kesennuma profile, Yoshii and Asano, 1972; Okada et al., 1979 ; see A–AVin Fig. 1).T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 55
T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 56
underlain an about 10-km-thick layer with a velocity of
6.2 to 6.3–6.45 km/s ( Iwasaki et al., 1994 ;Fig. 3(a) ).
Such a ‘‘middle crust’’ with an intermediate velocity of6.2–6.5 km/s is also found in Hokkaido (Iwasaki et al.,
1998) , central Japan (Takeda, 1997) and SW Japan
(Sakai et al., 1994) , probably representing a common
feature for the Japanese islands (Iwasaki et al., 2002) .
The seismic attenuation within the upper/middle crust
is large. Actually, Qp values from seismic wide-angledata in NE Japan and Hokkaido (Iwasaki et al., 1994,
1998) are 100–400 and 100–300 in a frequency range
of 5–15 Hz, respectively, showing a marked contrastwith the continental crust.
The lower crustal velocity obtained from the early
s t u d i e si s6 . 6 – 6 . 7k m / s( e . g . Hotta et al., 1964;
Hashizume et al., 1966; Yoshii and Asano, 1972;
Yoshii et al., 1974 ). These results, however, may
contain large estimation errors because the lowercrustal arrivals usually appear in a limited offsetrange. Amplitude analysis for the 1990 experimentin NE Japan showed that a relatively high velocity(6.9–7.0 km/s) layer is situated at the base of thelower crust ( Iwasaki et al., 1994 ;Fig. 3(a) ). However,
a velocity exceeding 7 km/s, as reported for the lower
crusts of shields, platforms or passive margins (Hol-
brook et al., 1992) , has not been reported except for
one case in central Japan (Takeda, 1997) .
Lower crustal wide-angle reflections are observed
in many regions of eastern Hokkaido (Iwasaki et al.,
1998) , NE Japan (Iwasaki et al., 1994, 2001) , central
Honshu (Takeda, 1997) , SW Japan (Hashizume et al.,
1981) and eastern Kyushu (Research Group for
Explosion Seismology 1999a,b) . Waveforms from
these reflectors are not pulsive, but have reverberationwith 1–2 s duration time (Fig. 3(b)) . Such a feature is
interpreted by a laminated lower crustal model com-posed of thin (0.3–1 km) alternating layers with low(6.0–6.5 km/s) and high velocity (6.5–7.0 km/s)velocities (Takeda, 1997, Iwasaki et al., 1998) .W e
think these features indicate ‘‘reflective lower crusts’’
as are widely observed in continental regions (e.g.Barazangi and Brown, 1986a,b; Mooney and Meiss-ner, 1992 , see Section 3.1).The Pn velocity under the Japanese islands is
estimated to be low (7.5–7.9 km/s) from rather oldexperiments. PmP phases observed in the recentexperiments (Fig. 3(b)) are sometimes characterized
as a wave train with a duration time of 2–3 s. Thisfeature indicates that the Moho forms a transition zonefrom the upper mantle to the lower crust without a
large velocity contrast.
3. New crustal images from recent
multidisciplinary projects
3.1. NE Japan Arc
Since 1997, we started a new multidisciplinary
project for elucidating deformation processes of islandarc crust through well-organized expedition of con-trolled source seismic survey, geological observationand microseismic observation. The first expedition forthis project was undertaken in NE Japan during1997–1998 ( Hirata et al., 1999a,b ;Matsubara, 2002 ;
Iwasaki et al., 2001 ;Sato et al., 2002 ;Fig. 1 ). An
onshore–offshore wide-angle seismic experiment in
1997 aimed at elucidating large-scale structural heter-ogeneity from the trench to the back arc basin (profile
‘b’ in Fig. 1) . The onshore part of the profile is 150.9
km in length, on which 287 temporal recordingsystems were deployed to observe six 500-kg andfour 100-kg explosive shots ( Iwasaki et al., 2001 ;Fig.
4(a)). It is noted that this profile crossed the 1990
profile line in the Kitakami Mountains (Figs. 1 and 3) .
Record sections obtained show travel time undula-tions of first arrivals west of the Kitakami Lowland,indicating a complex uppermost crustal structure (Fig.
4(b)) . The first arrivals within the Kitakami Moun-
tains, on the other hand, lie on a simple travel timecurve with an apparent velocity of 6 km/s. Severallater phases are not pulsive, but recognized as a wave
train probably generated within the reflective middle/
lower crust.
The crustal model constructed by a ray-tracing
technique clearly shows the deformation under the
Fig. 3. (a) Crustal structure model for the eastern part (the Kitakami Mountains) of NE Japan ( Iwasaki et al., 1994 , see ‘a’ in Fig. 1 ). Note a
middle crust with a velocity of 6.2 to 6.3–6.45 km/s. Lower part of the crust is highly reflective. Triangles at the top of the model show shotpoints. This section crosses the 1997 profile south of S-2. (b) Example of record section (shot S-1) obtained in the eastern part (the Kitakami
Mountains) of NE Japan. Note a number of wide-angle reflections from the middle/lower crust and uppermost mantle (see (a)).T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 57
Fig. 4. (a) Location map of the 1997 refraction/wide-angle reflection experiment in NE Japan (Iwasaki et al., 2001) . Stars indicate shot points of
500 (L-1–L-6) and 100 kg (M-1–M-4) charges. Solid lines indicate seismic reflection surveys undertaken in 1997–1998. Quaternary volcanicfront (Q.V .F.) is shown by a broken line. (b) Record section from shot M-3 (a). Note the prominent travel time undulation in the western side of the
shot and the reverberations of reflected waves in the eastern side. Solid curves show travel times predicted from a model in (c). R1–R4 indicate
reflections from midcrustal interfaces in (c). (c) Crustal structure model across NE Japan by the 1997 seismic refraction/wide-angle reflectionprofile ( Iwasaki et al., 2001 ; see ‘b’ in Fig. 1 ). Note the highly deformed sedimentary layers and prominent crustal thinning in the western part of
the profile. Hypocenters by dense seismic network are also plotted ( Matsubara, 2002 ). This section crosses the 1990 profile in the Kitakami
Mountains (see Figs. 1 and 3 ). a: Sen’ya Fault; b: Kawafune Fault; c: Uwandaira Fault (see (a) and (d)). Q.V .F.: Quaternary volcanic front. (d)
Crustal section (unmigrated) in the backbone range of NE Japan by seismic reflection survey (a). Note U-shaped fault geometry bounding thebackbone range. The Sen’ya Fault becomes horizontal at 4.7–5 s TWT below which a number of reflectors are developed.T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 58
Fig. 4 ( continued ).T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 59
extensional tectonics associated with the Miocene
back arc spreading (Fig. 4(c)) . Travel time curves
predicted from the model are superimposed on theobserved section (Fig. 4(b)) . They satisfactorily
explain both the first and later arrivals. The easternpart of the profile has been a stable forearc block sinceMiocene, which is composed of a less deformed
upper/middle crust of higher velocity (6.0 to 6.05–
6.3 km/s) and a number of reflectors below a depth of10–12 km (see also Figs. 3(a,b) and 4(b) ). The total
crustal thickness of this block is 32–33 km. Thewestern part of the profile is covered with thickMiocene sediments. The upper, middle and lowercrustal velocities are 5.75–5.85 to 5.90–6.15, 6.15–6.30 and 6.6–6.7 km/s, respectively. The crustal
thickness reaches its maximum near the present vol-
canic front, and keeps almost constant 10–15 km tofurther east. This may be explained by the highvolcanic activity (magmatic intrusion/underplating)in the last 10–15 Ma (Sato and Amano, 1991) .
Actually, the volcanic front before 8–10 Ma wassituated several 10 km east of the present position.The crustal thickness decreases to 27 km at the west-
ern end of the profile. This westward crustal thinning
begins at the eastern edge of the backbone range,almost coincident with the eastern limit of the Mio-cene normal faults and extensional basins. Probably,the thinning represents the crustal stretching duringthe Early Miocene.
As shown in Fig. 2 , the lateral variation in Pn
velocity is evident across the NE Japan Arc. The
recent ocean bottom seismographic observations also
support a velocity value of 7.9–8.0 km/s beneath thesea of Japan (Nishizawa and Asada, 1999; Nishisaka,
2000) . Land data from the 1997 experiment, on the
other hand, require a low Pn velocity of 7.6–7.7 km/sbeneath the western half of NE Japan. These resultsmean that the Pn velocity change occurs in a 20–30-km-wide transition zone beneath the western coast of
NE Japan. Physical explanation of this transition is
still enigmatic. Although the thermal regime plays animportant role, it cannot sufficiently explain such asharp transition zone.
Seismic reflection surveys were conducted in 1997
and 1998 for mapping fine structures of active faultsand midcrustal reflectors under the backbone range(Sato et al., 2002 ). The total profile length of these
surveys is 38.4 km. Seismic signals from 136 vibro-seis and 16 dynamite shots were acquired by a digital
telemetry system with 779 channels. The charge sizesof the dynamite shot were set to 50–500 kg. Majorfaults in NE Japan were originally formed under theMiocene crustal extension, and were reactivated as areverse fault by the compressional stress field startingat 3 Ma. Our 1997 profile line crosses the Sen’ya and
Kawafune faults in its western part (Fig. 4(a)) . These
faults are considered to be conjugate to each other,and responsible for the 1896 Riku-u earthquake(M7.2).
Fig. 4(d) shows a crustal section from the reflection
data, in which the major active faults under the back-bone range are clearly mapped down to a depth of12–15 km. The locations of these fault images are
also depicted in Fig. 4(c) . The area bounded by these
faults is 2–2.5 km uplifted with respect to the sur-rounding areas under the compressional stress regime(the pop-up structure). The travel time and amplitudeanalyses for the reflected data indicate that the Sen’yaFault is characterized by a 0.5–1-km-thick low-veloc-ity material (0.5 km/s lower than in the surroundedarea). This result is also confirmed by our wide-angle
data. Actually, there is no abrupt travel time jumps or
amplitude attenuation indicative of a large low veloc-ity body (more than 2–3 km thick) beneath thebackbone range. The fault plane becomes almosthorizontal at a depth of 12–13 km, beneath which anumber of reflectors are developed. The depth rangeof these reflectors is almost consistent with the reflec-tive zone deduced from wave trains (reverberation) of
the wide-angle reflections (see Fig. 4(b) and (c) ).
According to the hypocentral distribution by the
dense seismic networks ( Matsubara, 2002 ), most
events are concentrated in the upper 12–15-km crust(Fig. 4(c)) . These results strongly indicate difference
in rheological properties between the upper seismo-genic and lower reflective parts of the crust. We alsosee a good correlation between the fault geometry and
the seismic activity.
3.2. Hidaka Collision Zone, Hokkaido
The first deep seismic reflection surveys in Japan
were carried out from 1994 to 1997 in the southernpart of Hidaka Collision Zone, Hokkaido (Arita et al.,
1998; Tsumura et al., 1999; Ito, 2002) . The NE part of
the crustal section is characterized by two distinctT. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 60
reflective zones (Fig. 5) . Namely, the upper 23 km
part of the colliding Kuril Forearc is obducted west-ward onto the NE Japan Arc, while the lower part isdescending down. These images provide a directevidence on the ongoing delamination of the KurilForearc. The wedge of NE Japan Arc is intrudingeastward into the delaminated Kuril Forearc.
Seismic expeditions in 1998–2000 aimed at inves-
tigating the precise crustal structure and its deforma-tion style in the entire part of the Hidaka Collision
Zone. Two reflection profiles in 1998 and 1999 wereundertaken to map the deeper structure in the northernpart of the collision zone. These lines are totally 83.4km long, on which 1567 receivers were set with aspacing of 50–150 m to record 74 vibroseis and 30explosive shots ( Iwasaki et al., 2000a,b ; see the index
map in Fig. 5 ). The charge size of dynamite was
ranging from 40 to 500 kg. A 227-km refraction/wide-
Fig. 5. Migrated crustal section in the southern part of Hidaka Collision Zone (modified from Ito, 2002 ). Locations of the profiles are designated
by 94, 96 and 97 in the index map (Arita et al., 1998; Tsumura et al., 1999; Ito, 2002) together with another reflection lines in 1999–2000 in the
northern part of the collision zone (Iwasaki et al., 2001, 2002) . Two distinct dipping reflective zones in the western part of the section indicate
crustal delamination occurring within the Kuril Forearc.T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 61
Fig. 6. (a) Preliminary crustal section (migrated) in the northern part of the Hidaka Collision Zone (see Fig. 5 ). A series of eastward dipping events are mapped, which are interpreted as
midcrustal reflectors in the obducted Kuril Forearc. Nk: Nakanogawa Gr. Q3: Quaternary sedimentary rock, Hm: Hidaka metamorphic rock. (b) A crustal model of the Hidaka Collision
Zone from the wide-angle data (Moirya et al., 2001) . Geometry of midcrustal interfaces of R1 and R2 are well correlated with those deduced from the seismic reflection data (a).T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 62
angle line was shot in 1999 to elucidate large-scale
structural variation from the NE Japan Arc to the
Kuril Forearc ( Moirya et al., 2000 ;profile ‘c’ in Fig.
1). On this profile, 297 receivers were set and six
dynamite shots were detonated. The central part of theline was coincided with the abovementioned reflec-tion profiles.
In a migrated section of the reflection data (Fig.
6(a)) , we see two events (R1 and R2) dipping east-
ward at 2–10 s TWT. Fig. 6(b) shows an upper crustal
model of the collision zone determined from the wide-
angle data (Moirya et al., 2000) . The structure east of
the Hidaka Mountains is characterized by highlyundulated sedimentary layers and a relatively low(5.6 km/s) crystalline basement. Beneath the HidakaMountains, the basement almost outcrops with ahigher velocity (5.9–6.1 km/s) than further east. Inthe region of this high velocity body, metamorphic
rocks of the middle/lower crust are exposed (Komatsu
et al., 1986) . The most important feature of the model
is the existence of two remarkable reflectors with aneastward dip at depths of 15–22 and 17–26 km.These reflectors are well correlated with the eventsimaged from the reflection data. Namely, the upperone represents the eastward extension of the event ofR1 in Fig. 6(a) , and the lower one almost coincides
with ‘‘R2’’. These results strongly indicate the west-
ward obduction of the middle/lower crustal materialsoriginally forming the Kuril Forearc. From these data,
however, the downgoing lower crust as in Fig. 5 has
not been imaged. At the moment, it is not clear
whether this reflects a difference in deformation stylebetween the northern and southern parts of the colli-sion zones. More integrated analysis of the refractionand reflection data are inevitably important to eluci-date the deeper part of the crustal deformation asso-ciated with the collision.
4. Discussion and conclusions
Seismic crustal studies with controlled seismic
sources in Japan have been providing importantinformation not only on physical properties of islandarc crust, but also on its evolution and deformationprocesses. The upper crust shows a significant heter-
ogeneity with a wide velocity range of 5.5–6.1 km/s.
Lateral variation of the velocity often occurs acrossmajor tectonic lines and/or geological boundaries.The lower crustal velocity is ranging from 6.6 to7.0 km/s. It is quite rare that it exceeds 7 km/s, asreported in a shield, platform or passive-marginregion (Holbrook et al., 1992) . The recent expeditions
often indicate the existence of a middle crust with an
intermediate velocity of 6.2–6.5 km/s in a depth
range of 5–15 km.
Fig. 6 ( continued ).T. Iwasaki et al. / Tectonophysics 355 (2002) 53–66 63
Deep crustal reflectivity is reported in many refrac-
tion/wide-angle reflection profiles (Iwasaki et al.,
1994, 1998, 2001, 2002; Takeda 1997; ResearchGroup for Explosion Seismology, Japan 1999a,b) .
Usually, it is recognized as wide-angle reflectionswith long duration time (reverberation). The recentseismic expedition in NE Japan which shows a
reflective zone deduced from such reverberations is
almost consistent with the deep crustal image fromseismic reflection data. The reflective part sometimesstarts in the lower part of the middle crust at a depth of10–12 km. Such a reflective property is in a closerelation with the rheological structure of the crust.Actually, most of shallow earthquakes are concen-trated within the ‘‘transparent’’ upper/middle crust,
which is in contrast with the low seismicity in the
‘‘reflective’’ lower crust (Ito, 1993, 1999) . Although
seismic activity within the middle crust is not eluci-dated with sufficient accuracy at the moment, themiddle crust may correspond to a transition zone froma brittle to a ductile rheology.
The origin of the high reflectivity in the deep crust
is not clarified as yet. As possible reasons, Mooney
and Meissner (1992) pointed out the magmatic intru-
sion and the metamorphic shearing. The former originis plausible for the case of central Japan because thelamination pattern is found in a thermally activeregion. Ito (1993) also pointed out the correspondence
between the heat flow and the cut-off depth of theshallow seismic activity. Another possible explanationis fluid within the crust. In the eastern part of NE
Japan, local reflectors are situated at the upper boun-
dary of the high conductivity zone which was formedby trapped fluid (Ogawa, 1992) .
The uppermost mantle beneath the Japanese islands
is characterized by a low Pn velocity of 7.5–7.9 km/s.Observations on rather weak PmP phase suggests thatthe Moho is not a sharp boundary with a large velocityjump, but forms a transition zone.
The seismic reflection studies since 1994 suc-
ceeded in mapping detailed deep structures of activefaults in NE Japan. The fault images obtained showlistric fault geometry, which becomes almost flat at adepth of 12–13 km where a reflective zone starts.This may also indicate the difference in rheologicalproperties between upper and lower parts of the crustas described above. The expeditions in Hokkaido
revealed the deformation style of arc–arc collision.A series of seismic refraction surveys imaged clear
crustal delamination of Kuril Forearc colliding to theNE Japan Arc. This is also supported by layergeometry and lateral velocity variation derived fromthe seismic refraction/wide-angle reflection experi-ments. Such a delamination process with aid of platesubduction at deeper depths plays a key role in
producing more felsic continental crust from an island
arc crust (Ito, 2002) .
The abovementioned velocity structures are close
to those of continental crusts (Christensen and
Mooney, 1995) , showing a marked contrast with
other arcs of an oceanic origin such as the Izu-BoninArc (Suyehiro et al., 1996) , the Aleutian Arc (Hol-
brook et al., 1999) and the Costa Rican Isthmus
(Sallares et al., 2001) .I nt h eA l e u t i a na r c ,f o r
example, the lower crust is characterized by highseismic activity and little reflectivity, indicating itscompletely different rheology from the case of theJapanese islands. The continental properties of theJapanese island crusts may have been attained at theeastern margin of the Asian Continent before theMiocene back arc spreading. Actually, the crust of
the eastern part of NE Japan, which has been a stable
block after the back arc spreading, has a reflectivezone with no seismicity in its lower part. Regionaldifference in arc magmatic processes may be anotherfactor for the structural variation in lower crust(Holbrook et al., 1999) .
Acknowledgements
The authors wish to express their sincere thanks to
Professor Hans Tybo for offering an opportunity tosubmit our manuscript. They also thank ProfessorDavid Okaya, University of Southern California, forhis critical reading of the manuscript and his valuablesuggestions.
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Lithos 184-187 (2014) 346-360 Contents lists available at ScienceDirect Lithos ELSEVIER Hitokabe Metasomatic hydration of the Oeyama forearc peridotites: rossMark Tectonic implications Toshio Nozaka * Department of Earth Sciences Okayama University, Okayama 700-8530,Japan ARTICＬE INＦO ABSTRACT Article history: In contrast to the widely recognized aspects of serpentinization, initial stages of hydration and tectonic processes Received 18 April 2013 of unserpentinized peridotites are still unclear, but have important implications for understanding the litho- Accepted 20 November 2013 spheric architecture of supra-subduction zones. This study provides petrological evidence from the Oeyama Available online 1 December 2013 and ductile deformation of forearc peridotites. Key findings in this study are: 1) complex association of high- Keywords: temperature metasomatic minerals: tremolitic amphibole, cummingtonite, phlogopite, chlorite, olivine and Exhumation Forearc peridotite orthopyroxene in veins and in mylonites; 2) the systematic variation in Si and Na + K contents of the tremolitic Hydration amphibole, corresponding to its mode of occurrence and mineral association; and 3) the presence of thin Metasomatism (<0.7 mm) veins of fine-grained olivine accompanied by a narrow diffusion zone of the host primary olivine. Mylonite On the basis of petrography and mineral chemistry, the temporal sequence of hydration and deformation of the Oeyama ophiolite is considered as follows: 1) infiltration of slab-derived fluids, causing decomposition of pri- mary pyroxene and chemical modification of primary olivine, 2) metasomatic formation of variable modal amounts of amphibole, phlogopite, chlorite, vein-forming olivine and secondary orthopyroxene at 650-750 °C; s ( s rs p i sd s no s ( comparable temperature conditions for metasomatism and early-stage mylonitization. Mylonitization occurred exclusively in hydrous peridotites, and the peridotite mylonites were preferentially overprinted by syntectonic serpentinization. Diffusion profles of olivine cut by a vein suggest rapid cooling immediately after the metaso- matic fluid infiltration. From these observations and calculations, it is concluded that the exhumation of the forearc peridotites was closely related to the infiltration of high-temperature metasomatic fluids and hydration occurred under a wide range of temperature conditions. @ 2013 Elsevier B.V. All rights reserved. 1. Introduction is common in the supra-subduction zone mantle (e.g., Franz et al., 2002; Ishimaru and Arai, 2011; Khedr and Arai, 2010; Marocchi et al., Geological, geochemical and geophysical studies have revealed that 2007; Mclnnes et al., 2001; Nozaka, 2005). Because the evidence of dehydration of a subducting slab and coupled hydration of the overlying Hondera mantle wedge are common in subduction zones and cause fundamental cally deformed peridotites, a linkage between fluid/melt flow and the geologic proceses such as arc magmatism, dehydration-induced earth- formation of shear zones has been pointed out (e.g., Arai et al., 2004; quakes and chemical element recycling (e.g., Hacker et al., 2003; Downes, 1990; Kelemen and Dick, 1995; Nozaka, 2005). However, the Hasegawa and Nakajima, 2004; Hattori and Guillot, 2003; Iwamori, entire sequence of alteration and tectonic processes in the supra- BJ-13-106 2012; Scambelluri et al., 2004; Schmidt and Poli, 1998; Stern, 2002; posed exposures of peridotite mylonites and schistose serpentinites Tatsumi et al., 1983; Wyllie and Sekine, 1982). In particular, it is well have been reported from the Alps (Hermann et al., 200o; Li et al., BJ-13-107 2004; Scambelluri et al., 1995), the serpentinites seem to have been tle rocks and plays an important role in the tectonic exhumation of formed by prograde metamorphism of altered oceanic peridotites and Tabashine BJ-13-101 Maekawa et al., 2001). Also there is a growing body of evidence from and later formation of the schistose serpentinites. mantle xenoliths and ophiolites showing that metasomatic alteration An important example of such linkage is the Happo ultramafic at higher-temperature (T) than that at which serpentinization occurs complex juxtaposed by a high-pressure (P) metamorphic belt of central Japan. Nozaka (2005) has shown that mylonitic shear zones within the * Tel.: +81 86 251 7883;fax: +81 86 251 7895. Happo peridotites record multiple events of fluid-related deformation E-mail address: nozaka@cc.okayama-u.ac.jp. and metamorphism: a first stage of high-T alteration (700-800 °C) 0024-4937/$ - see front matter 2013 Elsevier B.V. Allrights reserved. http://dx.doi.org/10.1016/j.lithos.2013.11.012 T.Nozaka/Lithos 184-187(2014) 346-360 347 Was overprinted by serpentinization at 400-600 °C during exhumation metamorphic rocks (Arai, 1980; Nishimura, 1998; Nozaka, 1999; of this ultramafic complex. More recently, Nozaka and Ito (2011) argued Nozaka and Shibata, 1994; Takeuchi, 2002; Tsujimori and Itaya, 1999; that during a sequence of hydration and exhumation, the formation of Tsujimori et al., 2000). Geological and geochemical characteristics of cleavable olivine took place in the Oeyama ophiolite, which is believed the Tari-Misaka and Happo ultramafic complexes are suggestive of to belong to an ophiolite belt that includes the Happo complex (Fig. 1). sub-arc or forearc mantle origin (Arai and Yurimoto, 1995; Khedr and Initial studies of the Oeyama ophiolite point to a widespread distri- Arai, 2010); consequently the Oeyama ophiolite could originate in a bution of minerals indicative of amphibolite-facies or higher-grade re- supra-subduction zone as well, although the peridotites of the Oeyama equilibration. This suggests that the ophiolitic complex was subjected ophiolite show less degree of depletion (Ishiwatari and Tsujimori, to pervasive high-T alteration before juxtaposition with surrounding 2003). blueschist-facies metamorphic rocks. A detailed study of the high-T The ultramafic complexes of the Oeyama ophiolite are in fault con- secondary minerals provides a key to unravel the linkage between alter- tact with Paleozoic formations and the Renge Belt high-P/T metamor- ation and tectonic processes in the ophiolite. This paper reports textural phic rocks, and have intrusions of Cretaceous or Paleogene granitic and chemical variations of secondary minerals in the peridotites of the rocks, and are covered by younger sediments or volcanics (Fig. 1b-d; Oeyama ophiolite in order to elucidate the spatial and temporal rela- Igi and Kuroda, 1965; Igi et al., 1996; Is0zaki et al., 2010; Kurokawa, tionships between alteration and deformation processes. The data sug- 1985; Uda, 1984; Uemura et al., 1979). Typical high-P/T metamorphic gest that high-T metasomatic hydration was closely associated with the minerals such as lawsonite and glaucophane have been reported from initiation of exhumation of forearc mantle peridotites. the Renge Belt metamorphic rocks in proximity to the ultramafic com- plexes (Hashimoto and Igi, 1970). The granitic intrusions formed con- 2. Geological setting and general description of the tact aureoles in the Oeyama and Wakasa complexes and surrounding Oeyama ophiolite rocks (Kurokawa, 1985; Nozaka and Ito, 2011; Uda, 1984; Uemura et al., 1979). The Oeyama ophiolite is a collective name of the Oeyama, The main components of the Oeyama ophiolite are serpentinized pe- Sekinomiya, Wakasa and other small ultramafic complexes exposed in ridotites, which also include tectonic blocks or intrusions of other rock the Renge high-P/T metamorphic belt of SW Japan (Fig. 1a; Ishiwatari, types: pyroxenite; gabbro; amphibolite; and jadeitite (Chihara, 1989; 1989, 1990; Isozaki et al., 2010; Nishimura, 1998). Gabbroic rocks and Igi and Kuroda, 1965; Kuroda et al., 1976; Kurokawa, 1975, 1985; amphibolites included in the ultramafic complexes show K-Ar radio- Tsujimori and Liou, 2004; Uda, 1984; Uemura et al., 1979; Yamaguchi, metric ages of 470-400 Ma (Nishimura and Shibata, 1989; Tsujimori 1990). The predominant lithology of the peridotites varies within the et al., 2000). The Oeyama ophiolite is similar to the Tari-Misaka, Oeyama ophiolite. The main lithology of the Oeyama and Wakasa com- Ashidachi, Ohsayama and Happo ultramafic complexes (Fig. 1a) in li- plexes is dunite, a considerable part of which seems to be cumulates thology, age of amphibolite blocks and juxtaposition with high-P/T from basaltic melts (Kurokawa, 1985; Nozaka and Ito, 2011), whereas TM AS OS Ultramaficcomplex WSFig.(c) 100km Fig. (b) SuoBelt(230-160Ma) Sangun high-P/T SK OE Renge Belt(330-280 Ma)- HP b 134°42E 35°22 5km Granitic rocks Mesozoic-Cenozoic sediments/volcanics Locality of tremolitic amphibole Pelitic/maficschists Paleozoic sediments/volcanics +fine-grained granoblastic Peridotite/serpentiniteAmphibolite/pyroxenite/gabbroic rocks other types Tremolite-in isograd of metaserpentinite by contact metamorphism Tr-in.- 35°27 135°13'E Fig. 1. (a) Distribution ofultramafic complexes and the Sangun high-P/T metamorphic belt (Renge and Suo Belts) in SWJapan (Ishiwatari, 1989, 1990; Isozaki et al.,2010; Nishimura, 1998; Takeuchi, 2002). Abreviations forultramafic complexes: OE, Oeyama; SK,ekinomiya; WS,Wakasa; OS, Ohsayama;AS,Ashidachi; TM,TariMisaka; HP, Happo. (b) Geological sketch map of the Sekinomiya area (Igi et al, 1996). (c) Geological sketch map of the Wakasa area (Uemura et al., 1979). (d) Geological sketch map of the Oeyama area with the Tr (tremolite)-in isograds in the contact aureole of metaserpentinites (Igi and Kuroda, 1965; Kurokawa, 1985; Nozaka and Ito, 2011; Uda, 1984). Localities of tremolitic amphibole examined in this study are also shown. 348 T. Nozaka/Lithos 184-187(2014) 346-360 the Sekinomiya complex is mainly residual harzburgite left after partial dominantly lizardite/chrysotile statically replacing olivine and melting (Arai, 1980). Most peridotites were serpentinized and de- orthopyroxene; 6) contact metamorphic minerals, including fine- formed to a variable extent (Fig. 2). Serpentinite mylonites character- grained granoblastic, highly magnesian olivine, antigorite, tremolite, ized by lepidoblastic antigorite, with or without porphyroclastic and talc and/or orthopyroxene (Fig. 3h, i); and 7) products of static re- neoblastic olivine, occur in the Oeyama ophiolite, as in the Happo com- hydration such as lizardite/chrysotile, talc and clay minerals. plex (Nozaka, 2005). The olivine porphyroclasts and neoblasts of the Tremolitic amphibole, an amphibolite-facies mineral in meta- Happo and Oeyama peridotites originated from high-T mylonitization peridotites (>450 °C; e.g., Bucher and Frey, 2002), is widely distributed of peridotite before the formation of serpentinite mylonites (Nozaka, in the Oeyama ophiolite, even in rocks unaffected by contact metamor- 2005; Nozaka and Ito, 2011). A sequence of hydration and deformation phism (Fig. 1b-d), suggesting a high-T metamorphic event older than probably caused the formation of cleavable olivine in the peridotites the peridotite serpentinization and the blueschist-facies regional meta- (Nozaka and Ito, 2011). morphism of surrounding rocks. Tremolite is a second-stage mineral of the above classification, and the minerals of this stage are the main tar- get of this study. Detailed descriptions of minerals of the other stages 3. Petrography have been carried out in previous studies (Arai, 1980; Kuroda and Shimoda, 1967; Kurokawa, 1985; Nozaka and Ito, 2011; Uda, 1984); 3.1. Sequence of metamorphic assemblages s n necessary. Samples collected from the Oeyama, Sekinomiya and Wakasa com- plexes were examined under the optical microscope. Based on textural observations such as grain size, mutual grain contact, and cross-cutting 3.2.Descriptions of high-T alteration minerals relationship, seven groups of different mineral assemblages and/or tex- tures have been recognized: 1) primary peridotite assemblage (Fig. 3a), 3.2.1. Tremolitic amphibole including coarse-grained olivine, orthopyroxene and clinopyroxene In addition to contact metamorphic tremolite showing fine-grained with exsolution lamellae, and chromite or Cr-spinel with equant shape granoblastic texture with metamorphic olivine (Fig. 3h; Nozaka and Ito, in dunite and vermicular shape near pyroxene in harzburgite; 2) second- 2011), tremolitic amphibole occurs in four modes in the Oeyama ary minerals and mineral aggregates suggestive of amphibolite-facies ophiolite: 1) fringes statically replacing primary clinopyroxene in alteration (Fig. 3b-e), including tremolitic amphibole, cummingtonite, vein-free harzburgites (Fig. 3a); 2) vein-like aggregates or pseudo- phlogopite, chlorite, poikiloblastic orthopyroxene and veins of fine- morphic aggregates after pyroxene (Fig. 3b-d); 3) inclusions in poikiloblastic orthopyroxene (Fig. 3e); and 4) porphyroclasts and deformation of the above two groups of minerals (Fig. 3f, g); 4) products neoblasts in mylonites (Fig. 3f, g). In thin section, the mode 2) tremolitic amphibole is most common. Different modes of tremolite, other than 2) ondary clinopyroxene (Fig. 3f, g); 5) products of low-T serpentinization, and 3), seldom occur in the same thin section, but frequently occur in C d 10 cm 10cm Fig. 2. Mesoscopic structure of serpentinized peridotites from the Oeyama ophiolite. (a) A typical outcrop of serpentinite mylonite with boudinaged metagabbro. (b) A hand specimen of typical serpentinite mylonite with a prominent foliation (sample # OE0837, Oeyama complex).(c) A hand specimen of gently deformed metasomatized harzburgite with a weak foliation. formation and metasomatism are almost absent in this specimen (sample # KW2307, Sekinomiya complex). T. Nozaka / Lithos 184-187 (2014) 346-360 349 samples from neighboring outcrops. However, the spatial relationship (Fig. 4), and are cut by serpentine but never cut it. Olivine in relatively Tono body, Kitakami plutons, is LREE depleted with a mild positive Eu thick veins coexists with phlogopite, suggesting its formation during a olitic amphibole of modes 2) and 4) coexist locally with phlogopite or metasomatic event. The olivine veins show a distribution pattern sug- cummingtonite. Undulatory extinction is common in these modes, but gesting their three-dimensional interconnection via hidden veins, and pattern with negative Eu anomalies characterizes most of the Abukuma do not cut but are connected with the pyroxene pseudomorphs de- fine-grained clinopyroxene (Al- and Cr-poor diopside, see Table 1) is scribed below (Fig. 4a-c). common, particularly in serpentinite mylonites. In rock samples in which the olivine veins occur, there are domains that consist of fine-grained aggregates of tremolite, chlorite, 3.2.2. Orthopyroxene decomposed phlogopite, secondary clinopyroxene, serpentine and a Extant orthopyroxene occurs less frequently than tremolitic amphi- older than 119 Ma (~130 Ma), and the adakitic, calc-alkaline, and bole, olivine and chlorite, but was more abundant in the protolith completely decomposed into dark brown pseudomorphs, which seem because serpentine pseudomorphs after orthopyroxene are not uncom- to be cryptocrystalline aggregates of calcic clay minerals and Fe- lated negative Nb and Ta anomalies relative to La. A LIL component oxides/hydroxides judged from optical and chemical characteristics. plutons from the main Abukuma zone (Fig. 2) are younger than 115 Ma The tremolite-bearing fine-grained domains are similar to primary py- suggesting primary and secondary origin. The primary orthopyroxene, roxene in size, spatial distribution and nearby existence of vermicular like the one commonly occurring in mantle peridotites of SW Japan spinel. These facts as well as modal composition suggest that the do- (Arai, 1980; Nozaka and Shibata, 1994, 1995), forms discrete grains mains are pseudomorphs after orthopyroxene with a subordinate and has exsolution lamellae of clinopyroxene (Fig. 3a). Secondary amount of clinopyroxene, although the fine-grained aggregates locally orthopyroxene forms aggregates or veins, and has no exsolution run off out of the pseudomorphs to form veins connecting the pseudo- lamellae, but shows poikiloblastic texture containing fine-grained trem- morphs or penetrating into grain boundaries of neighboring primary olite, chlorite and olivine (Fig. 3e). The secondary orthopyroxene is dif- olivine (Fig. 4c). ferent from contact metamorphic orthopyroxene in that the latter contains considerable minute grains of magnetite formed during 4. Analytical procedures serpentinization before the contact metamorphism (Fig. 3i). The chemical compositions of minerals were analyzed using an elec- 3.2.3. Cummingtonite tron probe micro-analyzer equipped with four wavelength-dispersive Cummingtonite occurs in association with tremolitic amphibole that spectrometers (JEOL JXA-8230) at Okayama University. Quantitative forms vein aggregates near primary or secondary orthopyroxene analyses were carried out with an accelerating voltage of 15 kV and a in a few samples. It forms epitactic overgrowth on some tremolitic am- probe current of 20 nA. Standards used were natural or synthetic oxides phibole crystals (Fig. 3c, d). Compared with tremolitic amphibole, and silicates. The applied matrix corrections followed the procedures of cummingtonite shows a slightly lower interference color and extinction Bence and Albee (1968), using alpha factors of Nakamura and Kushiro angle under the microscope. Discrete grains of cummingtonite have (1970). Representative analyses are listed in Table 1. never been found in this study. In many samples, there are serpentine Whole-rock compositions of selected samples were analyzed with pseudomorphs after extremely elongated amphibole crystals, most of an X-ray fluorescence spectrometer (Rigaku RIX 2000) at Shimane Uni- which have grown on tremolitic amphiboles. Although, in these sam- versity, following the procedures of Kimura and Yamada ( 1996). Analy- ples, the original amphiboles cannot be identified due to intensive ses are listed in Table 2. a r S4; latest Ordovician to earliest Silurian). The oldest fossil date from the 5. Chemical compositions of minerals the Tono body (Mikoshiba and Kanisawa, 2008). The calc-alkaline the Oeyama ophiolite could be masked by subsequent serpentinization. As mentioned above, the Oeyama and Wakasa complexes contain part of the western southern Kitakami zone was dated 113 Ma for the 3.2.4. Phlogopite by contact metamorphism. To avoid confusion due to the effects of mag- Phlogopite occurs as a constituent of microscopic veins in associa- matic differentiation and contact metamorphism, the folowing descrip- Hirahara et al. (2015). The Hikami granite samples are as enriched as tions are mainly based on the analyses of rocks from the Sekinomiya ent of pseudomorphs after pyroxene (see next Subsection 3.2.5). It also complex. The residual harzburgite therein is dominant and the effect occurs as elongated plates parallel to the foliation of mylonites (Fig. 3f). the negative Eu anomalies are clearly primary. A gabbroid from the Most phlogopite crystals are partly or completely decomposed to chlo- dikes. Also included in the description are the analyses of some rite, clay and opaque minerals. The decomposed phlogopite and clay harzburgites unaffected by contact metamorphism in the Oeyama minerals after phlogopite are optically and chemically different from ex- complex. zone west of the Hizume-Kesennuma fault show similar spatial-tem- d = = d p no on uoq = jo son udosqe ym 5.1. Primary spinel green, and have deficiencies of K20 and total oxides (~90 wt.%, see Table 1). The distribution of the pleochroic minerals suggests phlogopite was more abundant originally. chromian compositions in equilibrium with primary olivine and pyrox- 3.2.5. Olivine enes. The Cr-spinel from the Sekinomiya and Oeyama complexes has a Olivine shows variable modes of occurrence. Relatively large equant Cr/(Cr + Al) ratio of 0.27-0.59 and a Mg/(Mg + Fe2+) ratio of 0.37- grains are common in the primary assemblage (Fig. 3a). Small discrete Stewart, 2009) show lower abundances of trace elements as reflected in grains are included in poikiloblastic orthopyroxene (Fig. 3e). forearc peridotites, rather than abyssal peridotites (Dick and Bullen, Porphyroclasts with cleavage-like parting planes and fine-grained 1984; Ishi et al., 1992) as do those of the Happo, Tari-Misaka, Ohsayama neoblasts occur in mylonites (Fig. 3f, g). Granoblastic aggregates and Ashidachi complexes in the Renge metamorphic belt (Fig. 5). (Fig. 3h) and reticulate overgrowths (after serpentine) on relict primary olivine occur in contact aureoles (Nozaka and Ito, 2011). 5.2. Tremolitic amphibole Another type of olivine occurs as fine-grained aggregates that form microscopic veins (commonly <0.7 mm thick). The veins penetrate Calcic amphiboles with Si >7.0 and Na + K<1.0 atoms per formula s e d jo s o 350 T. Nozaka / Lithos 184-187(2014) 346-360 tremolite, and minor magnesiohornblende and edenite, according to association with the progress of hydration or alteration with decreasing the classification of Leake et al.(1997). temperature. On the other hand, prograde metamorphism after Tremolitic amphibole shows a systematic variation in alkali (Na and a n( + i) m m n e so o uidis K) and Si contents (Fig. 6). Nozaka (2005) has found that the ratio of tios, because of the high degree of depletion of alkali elements during (Na + K)/Si is a useful indicator of the origin of tremolitic amphibole. serpentinization (Nozaka, 2005). The variation in (Na + K)/Si ratio of It is common in metamorphosed peridotites and serpentinites that ret- retrograde amphiboles should be controlled by limited alkali contents rograde amphiboles show a broad range of alkali depletion in of primary calcic pyroxenes in mantle peridotites. Therefore, it is a b Dp T. Nozako / Lithos 184187 (2014) 346360 351 probable that amphiboles with much higher (Na + K)/Si ratios than host olivine. However, the decrease of Fo content also occurs at some usual retrograde amphiboles were formed by the addition of Na and distance from the vein contacts (Fig. 7a, b), suggesting the effect of hid- K, i.e., by alkali metasomatism (Khedr and Arai, 2010; Nozaka, 2005). den veins, which do not appear in the plane of the thin section, or the Most amphiboles from the Oeyama ophiolite, except for those of contact effect of submicroscopic veins observed as thin bright streaks in back- metamorphic origin, lie outside the prograde field and have an scattered electron images (Fig. 4h). The vein olivine, on the other (Na + K)/Si ratio higher than the group of retrograde tremolitic amphi- hand, does not show a systematic variation of Fo content with distance boles, suggesting a metasomatic origin of the amphiboles (Fig. 6). It is from veinhost contacts. Instead, relatively large grains have a composi- noteworthy that the tremolitic amphibole systematically changes com- tional zonal structure with Fo contents lower at rim than at core, and position in response to the species of coexisting mineral. Tremolitic am- smaller grains have still lower Fo contents than the large grain rims phiboles coexisting with phlogopite (Phl, Fig. 6) have relatively high (Fig. 4h). alkali contents. Tremolitic amphiboles included in poikiloblastic orthopyroxene (Opx w/o Cum, Fig. 6) have lower Si contents than those coexisting with cummingtonite (Cum w/o Opx,Fig. 6). In a 6. Discussion thin section containing both of cummingtonite and poikiloblastic orthopyroxene (Opx + Cum, Fig. 6), tremolitic amphiboles included 6.1. P-T conditions of alteration and deformation in the orthopyroxene tend to have lower Si content than those forming veins with cummingtonite. Most of porphyroclastic and neoblastic am- P-T pseudosections for several peridotites from the Oeyama phiboles in mylonites are similar in Si content to the group coexisting ophiolite in the system CaOFeOMgOAlO3SiO2HzO (CFMASH) with cummingtonite, but some of them are similar to the group were drawn using the software Perple_X ver. 6.6.8 (Connolly, 2005, coexisting with orthopyroxene (Fig. 6). 2009, updated 2013) with the dataset of Holland and Powell (1998, updated 2002) and the CORK equation of state of fluids (Holland and 5.3.Orthopyroxene Powell, 1991, 1998). Adopted solution models were Diener et al. (2007) or Dale et al. (2005) for clinoamphibole, Holland and Powell The secondary orthopyroxene is clearly different from the primary (1998) for olivine, Holland and Powell (1996) for orthopyroxene and orthopyroxene in chemical compositions (Table 1), particularly in Cr clinopyroxene, and Holland et al. ( 1998) for chlorite, and the ideal solu- tion models for anthophyllite, talc, antigorite and brucite. An example of and Ca contents, which are useful indicators for discrimination between primary and metamorphic orthopyroxenes from SW Japan (Nozaka, cummingtonite-orthopyroxene-bearing rock (sample # MK0608. 2010). The Cr2Os and CaO content of the primary orthopyroxene of Table 2) is shown in Fig. 8, which is very similar to the results for the Oeyama ophiolite is >0.6 wt.% and >0.5 wt.%, whereas that of the other rocks in Table 2. secondary orthopyroxene is <0.3 wt.% and <0.1 wt.%, respectively One of the most useful clues to the PT conditions is the presence of (Table 1). The secondary orthopyroxene is similar in chemical composi- Mg-Fe amphibole, i.e., anthophyllite or cummingtonite, because it oc- tion to contact metamorphic orthopyroxene from SW Japan, but it is curs under a limited range of P and T. In the CMSH system, anthophyllite clear that the secondary orthopyroxene of the Oeyama ophiolite de- is stable under relatively low P (<7 kbar) and a narrow range of T (600- scribed here occurs outside the contact aureoles. Furthermore, 700 °C).It has been noted, however,that the stability conditions for an- coexisting olivine and amphibole do not show chemical characteristics thophyllite expand into a wider range of P and T in the case of addition of those of contact metamorphic origin (Nozaka, 2003, 2011). Metamor- of iron to the system (Evans and Guggenheim, 1988; Katzir et al., 1999; phic orthopyroxene with low Cr and Ca contents has also been reported Nozaka, 2011; Ulmer and Trommsdorf, 1999). This general tendency is from the Happo complex (Khedr and Arai, 2010; Nozaka, 2005). confirmed here by the P-T pseudosections (Fig. 8). where MgFe am- phibole is stable at relatively high P conditions up to 1013 kbar, depending on amphibole solid-solution models. The T condition of 5.4. Olivine cummingtonite formation is roughly estimated to be 600-700 °C. If the cooling of the Oeyama ophiolite is a result of tectonic exhumation Primary olivine in Sekinomiya harzburgites have forsterite (Fo) con- (as discussed later, Section 6.3), then a relatively high P (e.g. 10- tents of 91.091.5 and NiO of ~0.4 wt.%, showing a similarity to olivine 13 kbar) and hence T of 650700 °C is plausible for cummingtonite in residual mantle peridotites from SW Japan (Arai, 1980; Nozaka, formation(Fig.8). 2010). Vein olivine tends to have Fo contents less than (though vari- Another indicator of T condition is orthopyroxene. Using the able) the primary olivine, and is clearly different from deserpentinized geothermometer of Al solubility in orthopyroxene(Witt-Eickschen metamorphic olivine, which contains abundant magnetite inclusion and Seck, 1991), equilibration T conditions of 700-750 °C and 1000- and has extremely high or low Fo contents (Nozaka, 2003, 2010). The 1140 °C were obtained for the secondary and primary orthopyroxene, vein olivine and the host olivine penetrated by the veins shows consid- respectively. The former is in agreement with the T range for the erable variations in Fo, NiO and MnO contents; they are much more var- secondary assemblage of orthopyroxene + olivine + tremolite + iable than ordinary primary olivine in residual mantle peridotites chlorite inferred from the PT pseudosections (Fig. 8). That is, the T (Fig. 7ac). Because the Fo content of host olivine decreases in proxim- condition for metamorphic orthopyroxene is higher than that for ity to the veins, it is evident that the veins affect the compositions of the cummingtonite. Fig. 3. Mode of occurrence of tremolitic amphibole and associated minerals. Abbreviations for minerals: Chl, chlorite: Cpx, clinopyroxene;: Cum, cummingtonite; Ol, olivine; Opx, orthopyroxene; Phl, phlogopite or decomposed phlogopite: Srp, serpentine; Tr, tremolitic amphibole. (a) Photomicrograph of tremolitic amphibole statically replacing primary o e sae aoe n jo ueo oassdss xdoo pe aord m saxo pseudomorphs after pyroxene (crossed polars, sample # MK0611). (c) Photomicrograph of vein-forming tremolitic amphibole overgrown by cummingtonite (crossed polars, sample # MK0608). (d) CaKo X-ray map of a part of (c), obtained with an electron-probe microanalyzer. Cummingtonite (dark) is distinguishable from Ca-rich tremolitic amphibole (green). (e) Photomicrograph of a grain of a vein-like aggregate of orthopyroxene showing poikiloblastic texture with inclusions of olivine, tremolitic amphibole and chlorite (crossed polars, sam- ple # Mko904). The orthopyroxene lacks exsolution lamellae of clinopyroxene, unlike the primary orthopyroxene shown in (a). (f) Photomicrograph of elongated crystals of olivine, trem- olitic amphibole, phlogopite, chlorite and antigorite in a serpentinite mylonite (crossed polars, sample # NF0305). (g) Photomicrograph of porphyroclastic olivine with a tail composed of olivine, tremolitic amphibole and chlorite neoblasts in a serpentinite mylonite (crossed polars, sample # Y0613) (h) Photomicrograph offine-grained granoblastic olivine and tremolite in uoeueuoojo deoo ()( auessepod psso） xau ueojouoz on +e+uo one euawoj uudseau orthopyroxene, which contains abundlant tiny grains of magnetite, from the highest-grade aureole (olivine + orthopyroxene zone) of the Oeyama complex (crossed polars, sample # OE0830). 352 T. Nozaka/Lithos 184-187(2014) 346-360 Table 1 Representative microprobe analyses of minerals from the Sekinomiya complex of the Oeyama ophiolite. Mineral 01 01 01 01 01 01 01 Opx Opx Opx Opx Cpx Cpx Type/MO0 prim prim host(145）host(48）vein porphy neo prim prim poikilo poikilo prim prim Coexistingphase Phl Phl Tr Ⅱ Sample KW2307 KW1505 NF0307 NF0307 NF0307 NF0305 NF0305 KW2307 KW1505 MK0904 MK0608 KW2307 KW1505 SiO2 40.7 40.3 41.5 41.6 41.6 40.2 40.0 55.3 55.7 57.4 If such sinistral faulting is restored, the U-Pb ages young to the 53.6 52.8 TiO2 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.01 <0.02 <0.02 0.01 0.04 0.03 Al203 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 3.91 2.71 0.71 0.81 2.84 2.68 Cr203 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.87 0.80 0.17 0.23 0.72 1.14 FeOa 8.2 8.9 8.6 9.5 9.6 10.1 10.4 5.80 5.64 6.35 6.32 2.01 1.82 MnO 0.21 0.15 0.11 rvoir. Dashed lines: Hf evolution lines of 1000, 2000, 3000, and 4000 Ma 0.14 0.16 0.17 0.26 0.21 0.14 0.20 0.17 0.18 Nio 0.44 0.37 0.43 0.34 0.42 southwest Japan) may be consistent with this model. Japan, and the inner zone (Osozawa, 1994). 0.06 fault (Osozawa, 1994, 1997b). The radiolarian ages of cherts in accre- <0.04 Izanagi plate subduction (Deng et al., 2016; Fig. 14). The Kitakami and <0.04 <0.04 MgO 49.6 50.7 49.8 49.1 49.0 49.3 49.2 By using the younger (middle Cretaceous to Paleogene) cooling ages 34.3 35.1 35.3 In the Sambagawa zone, we suggested that the quartz eclogite may 17.0 Cao <0.01 The Tanagura tectonic line is the major sinistral fault bounding <0.01 0.01 0.01 <0.01 <0.01 0.66 0.71 0.08 0.10 24.4 25.4 Na20 n.d. n.d. n.d. n.d. n.d. n.d. n.d. <0.01 <0.01 <0.01 0.01 0.18 The T(trench)T(trench)R triple junction commonly envisioned is K20 n.d. n.d. n.d. n.d. n.d. n.d. n.d. <0.01 <0.01 0.01 <0.01 0.01 0.01 Total 99.1 100.4 100.4 100.6 100.7 100.1 100.1 100.9 100.1 99.9 99.2 100.7 101.1 Cations/O= 4 4 4 4 4 4 4 6 6 6 6 6 6 Si 1.00 0.98 1.01 1.01 1.01 0.99 0.99 1.89 1.92 1.98 1.96 1.93 1.91 Ti n.d. n.d. n.d. n.d. n.d. n.d. n.d. <0.01 Wakabayashi, 2017). The adakitic granitoids in the Kitakami zone may <0.01 <0.01 <0.01 <0.01 Al n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.16 0.11 Jourmal of Asian Earth Sciences 184 (2019) 103968 0.03 0.12 0.11 n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.02 0.02 <0.01 0.01 0.02 0.03 Fe 0.17 0.18 0.17 0.19 0.20 0.21 0.21 0.17 0.16 0.18 0.18 0.06 0.06 Mn <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 0.01 0.01 <0.01 0.01 0.01 0.01 Ni 0.01 0.01 0.01 0.01 0.01 0.01 0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 Mg 1.82 1.84 1.80 1.78 1.77 1.81 1.80 1.74 pyroclastic or continental detrital material (Okamoto et al., 2004). have formed during such an event. 1.83 0.90 0.92 Ca <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 0.02 0.03 <0.01 <0.01 0.94 Within northeast Japan, the Hatagawa and Furaba faults (Fig. 2) Na n.d. n.d. n.d. n.d. n.d. n.d. n.d. <0.01 <0.01 <0.01 <0.01 0.01 <0.01 K n.d. n.d. n.d. n.d. n.d. n.d. n.d. <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 Total 3.00 3.01 2.99 2.99 2.99 3.02 3.01 4.01 4.02 4.00 suggests that juvenile slab migrated from south to north, and this event grating TRT triple junction (Fig. 14) (e.g., Atwater, 1970; Furlong, Osozawa, 1994, 1997a, 1997b) proposed that the plate motion change XMg 0.915 0.911 0.912 0.902 0.901 0.897 0.894 0.913 0.915 0.908 0.909 0.937 0.943 a Total iron as FeO. n.d, not determined; XMg =Mg / (Mg + Fe). Mineral abbreviations: Atg, antigorite; Chl, chlorite; Cpx, linopyroxene; Cum, cummingtonite; Ol livine; Opx, orthopyroxene; Phl, phlogopite; Spl, spinel; Tr, tremolitic amphibole. Moo, mode of occurrence; prim, primary peridotite assemblage; host, host crystal penetrated by vein with the dis- tance from vein-contact (micron) in parentheses; porphy, porphyroclast in mylonite; neo, neoblast in mylonite; poikilo,host crystal or inclusion showing poikiloblastic texture; sec, sec- minerals. The T estimate using cummingtonite and orthopyroxene is consis- Ito, 2011). The similarities in mineral assemblage, mineral chemistry tent with tremolite compositions. The Si content (or Al content and deformation style between the Oeyama ophiolite and the Happo substituting Si) of tremolitic amphibole that coexists with an Al- complex suggest a similar overall history of hydration and deformation buffering phase such as chlorite is a potential indicator of metamorphic with changing P-T conditions. Khedr and Arai (2010) have considered T conditions for peridotites (Evans, 1982; Nozaka, 2005). In the Oeyama the P of hydration of the Happo peridotites to be >16 kbar on the ophiolite, tremolitic amphibole coexisting with orthopyroxene and basis of the lack of talc. However, talc is also very rare in the Oeyama chlorite has lower Si contents than that with cummingtonite and chlo- ophiolite, which contains cummingtonite that has a lower limit of P rite (Fig. 6). Some samples contain two types of tremolite, one of than talc + olivine (Fig. 8). The formation of talc, which is stable which is included in poikiloblastic orthopyroxene, and the other is over- under relatively high silica activity, may require elevation of the silica grown by cummingtonite. Because cummingtonite formation at a later activity of fluids. In addition, the formation of Mg-Fe amphibole could stage and lower-T condition than orthopyroxene is suggested by texture be controlled by nucleation kinetics rather than by P and T conditions and T estimate, the two types of tremolite probably formed during a (Nozaka, 2011). Therefore, the alteration of the Happo complex, even period of alteration associated with cooling. though Mg-Fe amphibole and talc appear to be absent, did not neces- In peridotite mylonites or serpentinite mylonites from the Oeyama sarily take place at a higher P than the Oeyama ophiolite. ophiolite, tremolite porphyroclasts and neoblasts have Si contents rang- The temporal sequence and T conditions of metamorphism and de- ing from 7.6 to 7.9 apfu across the value of the boundary between formation of the Oeyama ophiolite are summarized in Table 3. orthopyroxene- and cummingtonite-associated tremolitic amphiboles (Fig. 6). This suggests similar T conditions for the mylonitization as for the orthopyroxene- or cummingtonite-forming alteration. 6.2.Modal and chemical variations of metasomatic minerals In the Happo ultramafic complex, which appears to form an ophiolite belt together with the Oeyama ophiolite (Fig. 1; Ishiwatari, The formation of phlogopite in veins undoubtedly resulted from 1989; Isozaki et al., 2010), Nozaka (2005） found two stages of metasomatism, i.e., K-enrichment by fluid infiltration, because the mylonitization at different T conditions: peridotite mylonitization at modal composition of rocks and chemical compositions of primary min- 700-800 °C and serpentinite mylonitization associated with high-T erals indicate deficiency in K in the original peridotites of the Oeyama serpentinization (retrograde antigorite formation） at 400-600 °C. ophiolite. The parallel enrichment of Na + K of tremolitic amphibole They were followed by low-T serpentinization (mainly static lizardite/ in phlogopite-bearing rocks (Fig. 6) provides evidence that the infiltrat- chrysotile formation of mesh and bastite textures or veins) at <300 °C ing fluids were rich in Na as well as K. The degree of this alkali metaso- (Evans, 2004). Comparable mylonitization and serpentinization stages have been presented for the Oeyama ophiolite as well (Nozaka and the site of alkali-rich fluid infiltration because Na + K contents of T.Nozaka/Lithos 184-187(2014)346-360 353 Cpx Spl Spl Tr Tr Tr Tr Tr Tr Ⅱ Cum Phl Phl Chl sec prim prim vein/psd vein/psd poikilo porphy neo repl-cpx gran vein vein decom vein Atg Phl Cum Opx Phl Phl 10 Tr 10 10 NF0305 KW2307 KW1505 NF0307 MK0611 MK0904 NF0305 KW2307 NF0313 MK0611 NF0307 NF0307 NF0307 55.2 <0.01 <0.01 57.5 57.6 55.1 56.5 56.3 52.4 58.4 58.3 43.5 39.0 32.6 0.01 0.01 0.02 0.04 0.02 0.07 0.04 0.09 0.04 0.03 <0.02 0.12 0.10 0.04 <0.02 40.6 Kay, R.W., 1978. Aleutian magnesian andesites: melts from subducted Pacific Ocean 0.65 0.67 2.82 1.37 1.25 5.46 0.06 0.09 11.4 8.73 12.9 0.07 27.9 40.9 0.33 0.13 0.40 0.07 0.15 0.59 <0.02 0.06 0.29 0.03 4.43 1.40 15.6 18.0 1.75 1.72 2.60 2.25 1.94 1.92 1.12 7.18 2.55 11.4 3.33 0.05 0.49 0.53 0.05 0.06 0.07 0.09 0.07 0.12 <0.02 0.25 <0.02 0.33 0.05 <0.04 0.16 0.06 <0.04 0.05 0.14 0.15 0.12 <0.04 0.06 0.06 0.17 0.14 0.20 17.7 16.2 12.5 23.8 24.4 Hirahara, Y., et al., 2015. Space variation of Sr-Nd-Hf isotopic compositions in from 23.5 23.7 23.9 Kano, K., Kurimoto, C., Iwaya, T, Hoshizumi, H., Matsuura, H., Makimoto, H., 2002. 30.8 28.2 27.5 34.7 25.8 <0.01 0.04 11.7 13.3 12.7 11.9 11.9 13.1 13.0 0.61 <0.01 Frost, B.R., Barnes, C.G., Collins, W.J., Arculus, R.J., Ellis, D.J., Frost, C.D., 2001. A <0.01 Jahn, B.M., Glikson, A.Y., Peucat, J.-J., Hickman, A.H., 1981. REE geochemistry and n.d. n.d. 2.47 0.37 0.96 1.59 1.95 1.11 0.02 0.11 1.25 0.22 0.04 <0.01 n.d. n.d. 0.08 0.02 0.04 0.12 0.07 0.02 0.02 0.03 7.73 2.90 <0.01 100.3 100.9 100.2 98.3 98.3 98.1 97.5 97.5 98.6 97.7 97.4 95.2 90.5 88.2 6 4 23 23 23 23 23 23 23 23 22 22 14 2.00 <0.01 ? <0.01 7.85 7.84 7.58 7.78 7.75 7.19 7.94 7.90 6.05 5.84 3.08 <0.01 <0.01 <0.01 <0.01 0.01 0.01 <0.01 0.01 <0.01 <0.01 <0.01 0.01 0.01 <0.01 <0.01 1.34 1.00 0.10 0.11 0.46 0.22 0.20 0.88 0.01 0.01 1.87 1.54 1.44 <0.01 0.62 0.98 0.04 0.01 0.04 0.01 0.02 0.06 <0.01 0.01 0.03 <0.01 0.33 0.04 0.37 0.45 0.20 0.20 0.30 0.26 0.22 0.22 0.13 0.81 0.30 1.42 0.26 <0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 <0.01 0.03 <0.01 0.04 <0.01 <0.01 <0.01 <0.01 <0.01 0.01 0.02 0.02 0.01 <0.01 0.01 0.01 0.02 0.02 0.01 0.96 0.68 0.56 4.84 4.94 4.75 4.83 4.87 4.88 5.06 6.22 5.84 6.15 4.89 1.00 <0.01 <0.01 1.71 1.94 1.87 1.75 1.76 1.93 1.90 0.09 <0.01 0.03 Ishihara, S., Hamano, K., Ikegami, A., 1998. Isotopic evaluation on the genesis of the 0.01 n.d. n.d. 0.65 0.10 0.25 Osozawa, S., Wakabayashi, J., 2012. Exhumation of Triassic HP-LT rocks by upright 0.52 0.30 0.01 0.03 0.34 0.06 0.01 <0.01 n.d. n.d. 0.01 <0.01 0.01 0.02 0.01 <0.01 <0.01 0.01 1.37 0.56 <0.01 4.01 3.02 3.00 15.41 Martin, H., 1999. The adakitic magmas: modern analogues of Archaean granitoids. Lithos Yan, J., Liu, J., Li, Q., Xing, G., Liu, X., Xie, J., Chu, X, Chen, Z., 2015. In situ zircon Hf-O 15.32 15.38 15.47 Tsuchiya, N., 2008. Petrogenesis of adakites and their geological significance. Earth Sci. 15.12 15.83 15.67 10.02 0.958 0.649 0.555 0.960 0.962 0.941 0.949 0.956 Ishihara, S., Orihashi, Y., 2015. Cretaceous granitoids and their zircon U-Pb ages across 0.976 0.884 Kubota, K, Takeshita, T., 2008. Paleocene large-scale normal faulting along the Median 0.812 0.949 tremolitic amphiboles are relatively low in phlogopite-free or vein-free olivine shows a tendency to decrease Fo content towards contacts rocks (Fig. 6). with veins. The overall correlation between the distance from vein con- Tremolitic amphibole that forms veins could have crystallized from tact and the Fo content of host olivine (Fig. 7a, b) looks to be a composite fluids with a Ca source outside. On the other hand, the exclusive occur- pattern of an intra-crystalline Mg-Fe diffusion, caused by the infiltration rence of tremolitic amphibole within or near the pseudomorphs after of relatively Fe-rich fluids, and the effect of either submicroscopic veins pyroxene (Fig. 4a, c), even in phlogopite-bearing, highly metasomatized (Fig. 4h) or hidden veins that do not occur at the thin-section surface. A rocks, suggests that Ca activity was controlled by the distribution of pri- high-temperature condition is favorable for the intra-crystalline diffu- mary pyroxene in rocks with a low water/rock ratio. In other words, un- sion of olivine by such microscopic-scale veining. like Na and K, Ca could be inherited from pyroxene (discrete or lamellar Like the hydrothermal olivine, metamorphic orthopyroxene forms clinopyroxene, or Ca-bearing orthopyroxene) in original peridotites. vein-like aggregates. Although there is some ambiguity because of Thus the infiltration of CaO-rich fluids into the forearc mantle, as sug- low occurrence, the formation of the vein-forming metamorphic gested by Ishimaru and Arai (2011), is not required for the amphibole orthopyroxene can be explained by either local variation of silica activ- formation in the Oeyama ophiolite. ity or variable degree of crystallization differentiation of metasomatic Vein-forming, fine-grained olivine seems to be of metasomatic ori- fluids. gin because of its coexistence with phlogopite (Fig. 4). Olivine is a Cummingtonite, which occurs still more infrequently, overgrows phase readily crystallized from hydrothermal fluids as shown by the on vein-forming tremolite grains in the proximity of primary or classic work of Bowen and Tuttle (1949). The thin, curvy veins of olivine secondary orthopyroxene. The presence of seeds of tremolite should o rd jo sq una e o su e o n favor cummingtonite formation because of its sluggish nucleation showing a distribution pattern suggestive of their three-dimensional in- kinetics (Bowen and Tuttle, 1949; Fyfe, 1962; Greenwood, 1963; Nozaka, 2011). The local elevation of silica activity near orthopyroxene morphs after pyroxene (Fig. 4c). It is probable that the vein-forming was probably another necessary condition for the cummingtonite olivine was crystallized, with or without phlogopite and chlorite, from formation. abundant domains with low silica activity. The composition of the 6.3. Compositional modification of olivine and tectonic implications fluid is obscure, but the low Fo contents of olivine within or near veins suggest a relatively Fe-rich composition of fluid. The fluid seems to It is obvious that metasomatic fluids that formed olivine veins have been differentiated by fractional crystallization of olivine because yielded the compositional heterogeneity of host olivine (Fig. 7). The the vein olivine varies in Fo content depending on grain size and has overall compositional variation of host olivine suggests that intra- compositional zoning (Fig. 4h). On the other hand, the host primary crystalline cation diffusion of olivine was caused by fluid infiltration at 354 T.Nozaka/Lithos184-187（2014)346-360 Fig. 4. Mode of occurrence of vein and host olivine.Abbreviations for minerals are the same as those of Fig. 3. (a) Whole thin-section scanned image (plane-polarized light, sample # NF0307).“Px psd"”means pseudomorph after pyroxene,which consists of tremolite and decomposed tremolite(dark brown),chlorite, decomposed phlogopite(pale brown),secondary clinopyroxene, serpentine and a small amount of talc.(b) Scanned image of the same area as (a) (crossed polars). (c) Sketch of the thin-section images of (a) and (b) showing the distri- bution of olivine ± phlogopite vein (red) and pseudomorph after pyroxene (blue). Olivine (gray) was shaded with a different darkness for the distinction of individual grains and grain boundaries. Serpentine and chromite are omitted in this figure except for several serpentine-rich domains (white). (d) Photomicrograph of a part of (b).The olivine + phlogopite vein is olivine ± phlogopite veins in a host olivine crystal in an extinction position (crossed polars). (f) Photomicrograph of a part of (d) (plane-polarized light). (g) Photomicrograph of the same area as (f) (crossed polars).Vein-forming olivine varies in grain size from 0.05 to 0.4 mm. (h) Back-scattered electron image of a part of (g). A relatively large grain of vein-forming olivine (yellow circle) has a compositionally zoned structure with forsterite (Fo) contents (red numbers) of 90.7-90.9 mol% at the core and 90.2-90.3 mol% at the rim. Smaller grains of olivine in the vein have lower Fo contents (89.7-90.0 mol%) than the large grain. Host olivine is darker than the vein olivine, but is brighter in proximity to the vein and has micron-scale bright streaks (red arrows), indicating a local increase of iron. high T. For the compositional heterogeneity to be preserved, on the variation of olivine and theoretical diffusion profiles provides a con- other hand, the fluid infiltration should be of limited duration, and the straint on the cooling rate of the Oeyama ophiolite (Fig. 9). host peridotites should be cooled to T low enough for diffusion to be Theoretical Mg-Fe inter-diffusion profiles of olivine were calculated sluggish or frozen. Therefore, comparison between the compositional on the basis of two simplified models: the first is diffusion from a source T.Nozaka/Lithos 184-187(2014)346-360 355 Table 2 Compositions of serpentinized metaperidotites from the Sekinomiya complex of the Oeyama ophiolite. Sample MK0608 MK0802 MK0611 MK0614 MK0613 MK0808 MK0828 MK0707 Amph/Px Tr, Cum, Opx Tr, Cum Tr, Cum Tr Ⅱ Tr SiO2 (wt.%) 45.34 45.76 46.04 44.07 43.90 46.00 47.95 44.93 TiO2 <0.01 <0.01 <0.01 <0.01 <0.01 0.02 <0.01 <0.01 Al2O3 1.67 1.04 0.99 1.51 1.15 1.35 1.67 1.22 Cr2O3 0.43 0.32 0.40 0.33 0.52 0.46 0.90 0.41 FeOa 7.14 7.55 7.33 7.00 8.21 6.91 6.40 7.19 Mno 0.12 0.12 0.12 0.11 0.12 0.12 0.14 0.12 Nio 0.29 0.32 0.30 0.29 0.30 0.28 0.27 0.30 MgO 41.38 42.61 41.62 42.71 43.20 41.78 39.16 42.97 Cao 1.74 0.08 0.67 0.69 0.72 0.98 0.10 1.25 Na20 <0.03 0'0> <0.03 0'0> <0.03 <0.03 <0.03 <0.03 K20 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 P205 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 <0.01 Total 98.11 97.80 97.47 96.71 98.12 97.90 96.59 98.39 LOI 11.02 12.40 11.56 11.73 11.18 10.14 12.02 15.08 Mg# 91.2 91.0 91.0 91.6 90.4 91.5 91.6 91.4 Total iron as FeO.LOI = loss on ignition. Mg# = 100 * MgO / (MgO + FeO) mole ratio.Amph/Px, metamorphic amphibole or pyroxene included; Cum, cummingtonite; Tr,tremo- litic amphibole; Opx, orthopyroxene. of convecting fluid, in which Fe concentration is maintained at a con- Jaoul et al. (1995). Similar vein-related variations in olivine composi- stant level; the second is diffusion after solidification of a sheet-like tions were observed in other samples as well, but the direction and vein. Each of the models is expressed by Eqs. (1) and (2), respectively shape of veins are not suitable for calculations. The initial Fo content (Crank, 1975; Shewmon, 1963): of the host olivine for the calculations was 91.15 mol%, which is the s o no 0.1 mm from vein contacts (Fig. 9). The initial composition of the diffu- C (1) (2√Dt/ sion source is unknown, and hence mg# [= 100 * Mg / (Mg + Fe)] of the source was tentatively assumed to be 89.5, considering the lowest erf (2) (2√Dt） (2√Dt/ O Vein/pseudomorphic aggregate where Co is the initial concentration of a component in the diffusion ● Porphyroclast/neoblast source, Cis the concentration at a distance x from the center of diffusion Poikiloblastic inclusion in Opx Static replacement of primary Cpx source after a time t, h is half the thickness of a sheet of diffusion source, 0.8 十Granoblastic(thermally metamorphic)aggregate and D is a diffusion coefficient. Phl The diffusion coefficients were derived from the experimental work 8 Metasomatic of Jaoul et al. (1995). The adopted h is 0.3 mm in each model. Olivine compositions used for calculations were from a representative sample 0.6- 8 O Opx + Cum (Figs. 4d and 7a), where a vein has linear contacts and the distance from the vein can be measured in the direction nearly parallel to [010] O 00 of host olivine. This is the same direction as in the experiments of (nd) y+ 1 de 1.0 ·OE,SK 0.4 RE 0 TM, AS, OS 口HP Opxw/oCum 0.2 M 0.0年 8.0 7.5 7.0 Si (pfu) Abyssal spinel perid Fig. 6. Na + Kvs. Si contents (per formula unit on the basis of 23 oxygen atoms) of trem- olitic amphiboles (Si > 7.0 and Na + K < 0.8) that coexist with chlorite. The fields for the 0.0 majority of tremolitic amphiboles of a different origin are shown: prograde amphiboles 0.0 0.5 1.0 (vertical rules; for data source, see Nozaka, 2005); retrograde amphiboles, excluding Mg/(Mg + Fe2+) those with vein-like,metasomaticmodes of occurrence (horizontal ules;fordata source, see Nozaka, 2005); and metasomatic amphiboles, excluding those statically replacing Fig.5. Cr/(Cr + Al) vs. Mg/(Mg + Fe2+) diagram for primary Cr-spinel (data source: Arai, clinopyroxene (shade; this study). Enclosed by solid lines are the compositions of metaso- 1980; Nozaka and Shibata, 1994; Khedr and Arai, 2010; Nozaka, unpublished data; and matic tremolitic amphiboles coexisting with phlogopite (Phl), with orthopyroxene and this study). The Fe2+ contents were calculated assuming spinel stoichiometry. Abbrevia- cummingtonite (Opx + Cum), with cummingtonite without orthopyroxene (Cum w/o tions for ultramafic complexes are the same as those of Fig.1.The compositional fields Opx), and with orthopyroxene without cummingtonite (Opx w/o Cum). The dashed line for spinels from Izu-Bonin-Mariana (IBM) forearc peridotites (Ishi et al, 1992) and abys- indicates an approximate boundary between tremolitic amphiboles coexisting with sal spinel peridotites (Dick and Bullen, 1984) are also shown for comparison. cummingtonite and those with orthopyroxene. 356 T. Nozaka/Lithos 184-187(2014) 346-360 a Vein Host 91.0 C o ○。° 8 jow) 90.5 品 8 90.0 ? 89.5 -0.2 -0.1 0.0 0.1 0.2 0.3 0.4 0.5 0.6 Distance from vein contact (mm) b 0.50 1SOH Vein Host C (%1M) 91.0 1% 。品 0.30 90.5 C ●Vein Ol O Host OI MnO 0.5mm 0.10 88 90.0 89.5 90.0 90.5 91.0 Line-profile distance Fo (mol%) Fig.7.Compositional variations of vein and host olivine. (a) Variation of forsterite (Fo) content of olivine with the distance from vein-host contacts.Allthe data were obtained from a vein and a single host olivine grain shown in Fig. 4d. The width of the vein is 0.6 mm, and its center is situated at -0.3 mm from the vein-host contact. In this and the following two figures, solid (red) and open circles indicate vein-forming olivine and host olivine, respectively. (b)A line profile of the Fo content of olivine across the vein of Fig. 4d.(c) Nio and MnO vs.Fo contents of the same vein and host olivine as shown in (a). Fo value of the vein and host olivine, although the actual fluid could be discussion that the compositional variation of the vein and host olivine less magnesian. was caused by high-T fluid infiltration and was frozen by subsequent A set of diffusion profiles (Fig. 9a) were calculated with Eq. (1) as- rapid cooling, which was probably associated with tectonic movements. suming prolonged infiltration of fluid at T of 800 °C and 700 °C, in The geological setting (juxtaposition with high-P/T metamorphic order to constrain the duration of fluid infiltration. Because the Fo con- rocks) and spinel composition (Fig. 5) of the Oeyama ophiolite suggest tent of host olivine decreases probably due to submicroscopic or hidden its forearc origin. The history of hydration, metasomatism and deforma- veins, the calculated diffusion profiles were compared with the group of tion of the Oeyama ophiolite (schematically illustrated in Fig. 10) takes highest Fo value at each distance from the vein. The values of the high- previous works on subduction zones (Hacker et al, 2003; Nozaka and Fo olivine were fitted by the calculated profile of t = c. 70 years at Ito, 2011; Peacock and Wang, 1999; Reynard et al., 2011; Schmidt and 800 °C and c. 300 years at 700 °C (Fig. 9a). Even though it is difficult Poli, 1998; Stern, 2002) into consideration. The chemical compositions to know the exact time required for diffusion, because of uncertainties of phlogopite and coexisting tremolitic amphibole suggest the metaso- of fluid composition and T condition, a shorter time (<100 years) is matic enrichment with K and Na, and the most likely origin of the meta- more likely, given that fluids had lower mg# than the assumed value somatic fluids is the dehydration of subducted slab (Fig. 10a). The and fluid infiltration took place at a higher T (>750 °C) than that of metasomatic fluids infiltrating the forearc mantle peridotites caused the crystallization of secondary minerals. Rapid cooling should follow the compositional modification (e.g., enrichment in Fe) of primary oliv- the fluid infiltration so that the diffusion profile was frozen. It is most ine and decomposition of primary pyroxene. It subsequently formed likely that such rapid cooling was caused by tectonic movement of the variable association of secondary minerals such as tremolitic amphibole, peridotite complex, associated with the development of ductile shear phlogopite, chlorite, olivine, orthopyroxene and cummingtonite at 650- Zones. The similarity in T between metasomatism and mylonitization 750 °℃, depending on the local variation of silica and alkali activities. (see Section 6.1) is in good agreement with this argument. The intensity of metasomatism changed gradually in the host perido- The peridotite complexes could be cooled at a great rate after the so- tites with patently metasomatized zones in proximity to veins lidification of fluids at lower T conditions as well. Another set of diffu- (Fig. 10b). Several factors suggest that it is likely that hydration- sion profiles (Fig. 9b), which were based on Eq. (2), assuming a induced softening (e.g., Drury et al., 1991; Hirth and Kohlstedt, 1996) promoted the tectonic movement of forearc mantle peridotites in asso- between solidified vein and host, or ferroan and magnesian olivine ciation with ductile deformation at localized shear zones (Fig. 10b, c). grains. During a period longer than ten to a hundred thousand years, Specifically: mylonitization is restricted to hydrous peridotites con- the intra-crystalline diffusion at 600-700 °C should make the solidified taining amphibole, phlogopite and chlorite (there are no anhydrous vein and host olivine more homogeneous than observed. Within a peridotite mylonites); metasomatic hydration and mylonitization of pe- shorter period, the complexes should be cooled to a lower T, at which ridotites occurred at similar T conditions; and olivine penetrated by diffusion is very sluggish. This would require a considerable drop of T veins has compositional variation suggestive of rapid cooling immedi- condition in a geologic time frame. It is concluded from the above ately after the fluid infiltration. Subsequently, mutually promoting T.Nozaka/Lithos 184-187(2014) 346-360 357 21 Table 3 a Ol Opx Chl Cpx the Oeyama ophiolite. Ol Atg 17— Brc Chl Metamorphic/deformation Main minerals Estimated T Cpx 17 stage Ol Atg 1. Equilibration in the Ol (primary), Opx (primary), >1000°℃ Chl Cpx upper mantle Cpx (primary), Spl 3 010px 2. Metasomatic hydration Tr, Cum, Phl, Chl, Opx (vein and 650->750°℃ poikiloblastic), Ol (vein and Chl Tr compositional modification of host of vein) 9F Ol Atg- 3. Mylonitization of hydrous Ol (porphyroclastic and neoblastic) 650-750°℃ Chl Tr Ol Opx Cpx peridotites ol TIc 4. High-T serpentinization Atg (lepidoblastic), Cpx (secondary) 400-600°℃ Ch Tr / and mylonitization 5F ol Ath 5. Low-T serpentinization Liz/Ctl (mesh, bastite or vein) 0.00> Chl Tr 6. Contact metamorphism Atg (granoblastic or interpenetrating), 300-750°℃ -0l Opx Cpx Tr Ol (granoblastic), Tr (granoblastic), 010pxTr Tlc, Opx (poikiloblastic) 400 500 600 700 800 900 7. Later low-T hydration Liz/Ctl (mesh, bastite or vein), 0.000> T(°C) Tlc, clay Mineral abbreviations: Atg, antigorite; Chl, chlorite; Cpx, clinopyroxene; Cum, 21 b cummingtonite; Liz/Ctl; lizardite or chrysotile; Ol, olivine; Opx, orthopyroxene; Phl, Ol Opx Chl Cpx phlogopite; Spl, spinel; Tlc,talc; Tr, tremolitic amphibole. Ol Atg 17 Brc Chl xd mineral assemblages and inhomogeneous chemical composition of Ol Atg 010px minerals resulted from variations in the composition of metasomatic Chl Cpx Chl Tr fluid and in the distance from the site of fluid infiltration. The Mg-Fe 3 diffusion profiles of olivine penetrated by veins of fine-grained olivine ± phlogopite ± chlorite suggest rapid cooling immediately 6 Ol Atg /ol Opx Cpx Chl Tr a Vein→ Host // ol TIc f ol Cum 91.0 OF Chl Tr Chl Tr -0l Opx Cpx Tr · Ol Opx Tr % 8. 90.5 400 500 600 700 800 900 ow) T(°C) 800°℃ 700°℃ Fig.8. P-T pseudosections for a representative cummingtonite-orthopyroxene-bearing 90.0 7x101yr 3x102yr 2 x10² yr 2 x 103 yr using the software Perple_X ver. 6.6.8 (Connolly, 2005, 2009, updated 2013) with the ..... 1 x 103 yr 5x103yr dataset of Holland and Powell (1998, updated 2002) and the CORK equation of state of fluids (Holland and Powell, 1991, 1998). Different solution models were adopted for am- -0.2 -0.1 0.0 0.1 0.2 0.3 0.4 0.5 phiboles: (a) Dale et al. (2005)for Ca-amphibole and the ideal solution model for antho- 0.6 phylite; and (b) Diener et al. (2007) for clinoamphibole. Other solid-solution models are Distance from vein contact (mm) the same between the two figures (see text in Section 6.1). The calculated compositions of clinoamphibole in these figures were categorized as tremolite or cummingtonite.Abbrevi- q ations for minerals: Atg, antigorite; Ath, anthophyllite; Brc, brucite; Chl, chlorite; Cpx, clinopyroxene; Cum, cummingtonite; Ol, olivine; Opx, orthopyroxene; Tlc, talc; Tr, 91.0 Q tremolite. hydration (high-T serpentinization) and deformation took place at % 90.5 lower-T, and serpentinite mylonites were overprinted in the shear jow) zones (Fig. 10d). The high-T metasomatic fluid infiltration could trigger the detachment of hydrated peridotites from the surrounding mantle 700℃ 0.009 wedge, and the subsequent rapid cooling associated with ductile defor- 90.0 5 x102 yr 1 x104 yr mation is a reflection of the tectonic exhumation of the hydrated peri- 5 ×103 yr 1 x 105 yr dotites to an area of partial serpentinization (Fig. 10a). The history of 5x104 yr - 5x105 yr the Oeyama ophiolite illustrates that hydration of peridotites under a 89.5 -0.3-0.2-0.10.0 0.5 0.1 0.2 0.6 wide range of T conditions plays an essential role in developing the Distance from vein contact (mm) architecture of supra-subduction zones. Fig. 9. Diffusion profiles that explain the compositional variation of vein and host olivine. 7. Conclusions Olivine compositions are the same as shown in Fig. 7a. Calculations for diffusion profiles were based on simplified diffusion models (Shewmon, 1963; Crank, 1975; see text in Peridotites of the Oeyama ophiolite originated in the forearc mantle Section 6.3 for details) and experimentally determined Mg-Fe inter-diffusion coefficients (Jaoul et al., 1995). (a) Profles with a diffusion source of convective fluids with mg# of and were subjected to infiltration by hydrothermal fluids. This resulted 89.5 at 800 °C (red lines) and 700 °C (blue lines). (b) Profiles with a diffusion source of in the metasomatic formation of tremolitic amphibole, cummingtonite, a solidified sheet-like vein of olivine with initial composition of Fos9.5 at 700 °C (red phlogopite, chlorite, olivine and orthopyroxene at 650-750 °C. Variable lines) and 600 °C (blue lines). 358 T.Nozaka/Lithos 184-187(2014)346-360 Arc crust Partial serpentinization C Exhumation Gently deformed 600℃ hydrous peridotite Mantle wedge 800℃ 10000℃- Asthenospheric. Hydrous peridotite flow Fig. (b, c) mylonite E-OI Gently deformed Slab s Fluidinfiltration hydrous peridotite Patentlymetasomatizedzone Exhumation, more hydration & (Phl + alkali-rich Tr) more deformation Vein Atg Fig.(c) Serpentinite Shearzone mylonite Host peridotite Increase of K, Na & Fe cross section of a subduction zone based on a synthesis of some previous works (Hacker et al, 2003; Peacock and Wang, 1999; Reynard et al., 2011; Schmidt and Poli, 1998; Stern, 2002). The thermal structure is that of SW Japan given by Peacock and Wang (1999) and Hacker et al. (2003), but the scales for vertical and horizontal distance are not shown because of the uncertainty of subduction angle for the Oeyama ophiolite.(b) Enlarged view of the site of mantle wedge metasomatism. Infiltration of metasomatic fluids forming olivine ± phlogopite, tremolite ± orthopyroxene, and tremolite ± cummingtonite veins. Formation of phlogopite and alkali-rich tremolitic amphibole around the veins depict an obvious metasomatic zone. Alkalicontents of tremolite andFe content of primary olivine of host peridotites increase toward theveins.Localized shearing tookplaceduring tectonic movement. (c)Enlargedviewof a part of (b). Peridotite mylonites with porphyroclastic olivine (P-ol and neoblastic olivine (N-Ol) with tremolite, and chlorite with or without phlogopite formed in the shear zones. The peridotite mylonites were surrounded by gently deformed massive peridotites composed mainly ofequant crystals of olivine (E-Ol)with variable degree of strain and dislocation density (Nozaka and Ito, 2011). (d) View of the same peridotite body as (c) after a certain time.Serpentinite mylonites formed by syntectonic high-T serpentinization, i.e., antigorite ± diopside formation under the stability conditions of olivine.This mylonitization was overprinted on the early-stage shear zones (i.e., peridotite mylonite zones）by mutual promotion of hydration and deformation during further cooling and exhumation (see also Nozaka, 2005). Cleavable olivine with well-developed parting planes was brought into prominence during the syntectonic serpentinization (Nozaka and Ito, 2011). after the fluid infiltration. Mylonitization of peridotites occurred exclu- Arai, S., Takada, S., Michibayashi, K, Kida, M., 2004.Petrology of peridotite xenoliths fm sively in hydrated peridotites, and the compositions of tremolitic am- Iraya volcano, Philippines, and its implication for dynamic mantle-wedge processes. Journal of Petrology 45, 369-389. phibole suggest a similarity in temperatures of the hydration and Bence, A.E., Albee, A.L., 1968. Empirical correction factors for the electron microanalysis of mylonitization. Later syntectonic serpentinization at 400-600 °C was silicates and oxides. Journal of Geology 76, 382-403. Bowen, N.L, Tuttle, O.F, 1949.The system MgO-SiO-HO. Bulletin of the Geological So- ciety of America 60, 439-460. these observations and calculations, it is concluded that the exhumation Bucher, K., Frey, M., 2002. Petrogenesis of Metamorphic Rocks, 7th ed. Springer-Verlag, of the forearc mantle peridotites was initiated in close association with Berlin. the infiltration of high-temperature metasomatic fluids and promoted Chihara, K., 1989. Tectonic significance of jadeitite in Hida Marginal Belt and Sangun Metamorphic Belt. Memoirs of the Geological Society of Japan 33, 37-51 (in by hydration under a wide range of temperature conditions. Japanese with English abstract). Connolly, JA.D., 2005. Computation of phase equilibria by linear programming: a tool for Acknowledgments geodynamic modeling and its application to subduction zone decarbonation. Earth and Planetary Science Letters 236, 524-541. Connolly, JA.D.,2009. The geodynamic equation of state: what and how.Geochemistry, I am grateful to Y. Sawada and R. Tayama for permission and assis- Geophysics, Ge0systems 10, Q10014. http://dx.doi.org/10.1029/2009GC002540. Crank, J., 1975. The Mathematics of Diffusion, 2nd ed. Oxford University Press, Oxford. tance for the analyses using the X-ray fluorescence spectrometer at Dale, J, Powell, ,White, R.W.,mer, F.L,olland, TJB,2005.A thermodynamic md Shimane University. My thanks also go to N. Fujiwara, Y. Ito, Y. Kawasaki for Ca-Na clinoamphiboles in Na2O-CaO-FeO-MgO-Al203-SiO2-H20-O for petro- and M. Komoda for assistance in field and laboratory works. The manu- script has been improved by valuable suggestions of P. Fryer, construc- Dick, HJ.B., Bullen, T., 1984. Chromian spinel as a petrogenetic indicator in abyssal and alpine-type peridotites and spatially associated lavas. Contributions to Mineralogy tive reviews of S. Arai and S. Endo, and comments from the editor, and Petrology 86, 54-76. M. Scambelluri. This study was supported by funding from the Wesco Diener, J.F.A, Powell, R, White, RW, Holland, T.J.B., 2007. A new thermodynamic model Scientific Promotion Foundation. for clino- and orthoamphiboles in the system NazO-CaO-FeO-MgO-Al2O3-SiO2- H20-0. Journal of Metamorphic Geology 25, 631-656. Downes,H., 1990.Shear zones in the upper mantlerelation between geochemical en- References richment and deformation in mantle peridotites. Geology 18, 374-377. Drury, M.R., Vissers, R.L.M, Van der Wal, D., Hoogerdujn Strating, E.H., 1991. Shear Arai, S., 1980. Dunite-harzburgite-chromitite complexes as refractory residue in the localisation in upper mantle peridotites. 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Earth and Planetary Science Letters, 62 (1983) 41-52 41
Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands
[31
Hydrogen isotopic compositions of hydrogen and methane from
some volcanic areas in northeastern Japan
Yasuhiro Kiyosu
Department of Earth Sciences, Faculty of Science, Nagoya University, Chikusa, Nagoya 464 (Japan)
Received April 21, 1982
Revised version received September 14, 1982
D/H ratios of volcanic condensates, hydrogen and methane collected from Hachimantai, Za~, Azuma, Nasu and
Kusatsu shirane volcanic areas in northeastern Japan were determined together with chemical components. On the basis
of the isotopic ratios for volcanic condensates, it is concluded that most of the water vapor is essentially local surface
water which has been heated and slightly enriched in 180 by exchange with silicates. 8D values of hydrogen and
methane range from -220 to -515%0 (SMOW) and from - 180 to -487%0 (SMOW), respectively. The D/H ratio in
volcanic hydrogen indicates that hydrogen may have isotopically equilibrated with water vapor at temperatures between
200 and 400°C. The 8D value of methane also suggests that volcanic methane has been in isotopic disequilibrium with
the water vapor and hydrogen and that most of the methanes may be thermogenic methanes produced by pyrolysis of
carbonaceous materials.
1. Introduction
Volcanic gases usually contain alkali-unab-
sorbed gases such as nitrogen, hydrogen, methane
and inert gases.
High-temperature (> 400°C) volcanic gas of
magmatic origin, which has not passed through a
groundwater body generally has high concentra-
tions of hydrogen. Presumably, volcanic hydrogen
is produced by the high-temperature reaction of
water with ferrous oxides and silicates. Whereas
data on hydrogen and oxygen isotopes in volcanic
condensates have been reported, only little is
known about the deuterium concentration of
volcanic hydrogen. Arnason and Sigurgeirsson [1]
analyzed the deuterium contents of the water vapor
and hydrogen from the oceanic volcano Surtsey
near Iceland and showed that total hydrogen con-
tained in the volcanic gases has 8D value of
-55.3~ vs. SMOW. Arnason [2] estimated the
bottom temperature of drill holes from the ob-
served D/H fractionation between hydrogen and water in geothermal areas of Iceland; the calcu-
lated temperatures were generally equal to or
greater than the observed temperatures.
The gas boiling from heated groundwater (<
400°C) in geothermal systems often contains more
methane than hydrogen. The concentration of
methane may be controlled by the chemical reac-
tion of carbon dioxide and hydrogen [3]. This
reaction equilibrium shifts to the side of enrich-
ment of methane as temperature decreases and
pressure increases. The methane observed in
volcanic gases may be created by this reaction,
and/or by the near-surface reaction of volcanic
gases and heat with organic material [4]. Very few
data have been reported on the variation of D/H
ratios in volcanic methane because this gas is
relatively uncommon as compared to components
like hydrogen.
In order to clarify the origins of volcanic hydro-
gen and methane, the present isotopic investiga-
tions were conducted. In some volcanic areas, an
attempt was also made to compare measured tem-
0012-821/83/0000-0000/$03.00 © 1983 Elsevier Scientific Publishing Company
42
peratures and isotopic temperatures calculated
from the isotopic fractionation between water
vapor and hydrogen.
2. Sample localities
Gas samples were collected from fumaroles of
Za~, Azuma (Issaikyo), Nasu and Kusatsu shirane
in 1980 to 1981, and from six dry-steam wells in
the Matsukawa geothermal area (Hachimantai) in
1977 and 1981. The sampling localities are shown
in Fig. 1.
Za~, ,4zuma and Nasu volcanoes
Za~, Azuma (Issaikyo) and Nasu are strato-
volcanos made of pyroxene andesite, belonging to
the northern subzone of the Nasu volcanic zone
[5]. The volcanic rocks overlie the Miocene base-
ment of green tuff, volcanics and sediments, and
pre-Tertiary granodiorite. The present activity
Hachimantai
(usatsu
e
0 1 O0 200 Krn
i i I
Fig. 1. Sampling localities. consists of acid hot springs and fumaroles con-
centrated around these volcanic areas. The Nasu
samples discussed here are gas samples collected
from Oana and Mugen fumaroles in this area.
Kusatsu shirane volcano
Kusatsu shirane, located at the intersection of
the extension of Nasu and Fuji volcanic zones, is a
stratovolcano built from pyroxene andesite [6,7].
The most recent activity was a minor steam erup-
tion which occurred around the pit of Mizugama
at the summit crater, in March 1976. There are
four principal fumaroles and acid hot spring areas:
Kusatsu shirane (the northern foot of the volcano),
Sesshogawara, Kusatsu yubatake and Manza
karabuki.
Matsukawa geothermal fieM
Matsukawa is northwest of Mt. Iwate in the
Hachimantai volcanic groups which belongs to the
Nasu volcanic zone. Natural activity at Matsukawa
consists only of hot springs. The area is one of
andesitic volcanism, and the volcanic activity is
divided into the Quaternary volcanic complex and
the lower Pleistocene Tamagawa welded tuff com-
plex, with the dacitic or andesitic tuff. Underlying
the Tamagawa complex is the Miocene Yamatsuda
formation, consisting of sandstone, siltstone and
tuff. The Tamagawa welded tuff is the main res-
ervoir for production wells. The hydrothermal al-
teration of the andesite has several zones with
increasing depth. At depth there are radially dis-
tributed zones of alteration consisting of chlorite,
montmorillonite, kaolinite and alunite. A relic al-
teration pattern of pyrophyllite is recognized in
the deeper wells [8-10].
3. Experimental
3.1. Sample collection
The sampling method was modified after that
of Mizutani [11] and Ozawa [12] for chemical and
isotopic analyses, as follows: Volcanic and geo-
thermal gas samples were collected in a water-
cooled KOH solution by introducing the gases
from a fumarole through a glass tube and from a
steam vent on the steam pipeline through rubber
tubing, respectively. The hydrogen, methane and
other alkali-unabsorbed gases were collected in a
pyrex gas collector with a volume of about 100 ml.
When the alkali-unabsorbed gases were collected
to a sufficient amount, both ends of the collector
were sealed off with a torch. The condensates, hot
spring waters and river waters were sampled for
isotopic analysis. The outlet temperature of the
fumarole was measured with a mercury thermome-
ter.
3.2. Analytical methods
Chemical composition of gas condensates in
alkaline solutions was analyzed by the method of
Ozawa [12], except for H 2 S determination. Carbon
dioxide was determined by the micro-diffusion
method. Sulfur dioxide and hydrogen sulfide were
determined gravimetrically as BaSO 4 and Ag2S,
respectively. The chloride was determined photo-
metrically with mercuric thiocyanate and iron
alum. The residual gases such as hydrogen, nitro-
gen, methane and argon were analyzed by gas
chromatography.
Deuterium analysis was carried out by passing
5 mg of water over hot uranium and comparing
the resulting hydrogen gas with a standard of
known deuterium content in a dual collector mass
spectrometer [13,14]. For 180 analysis, 2 ml of
water sample was equilibrated isotopically with 4
ml of CO 2 at 25°C [15]. This CO 2 was isotopically
analyzed with a mass spectrometer.
In order to determine the hydrogen isotopic
compositions of volcanic hydrogen and methane,
these gases must be separated from other alkali-
unabsorbed gases. The separation of hydrogen and
methane from other gas components was made by
a gas chromatograph on a 3 m molecular sieve 5 A
column with helium carrier gas [4]. After separa-
tion, the hydrogen was passed over a CuO fur-
nance, heated to about 500°C and the combustion
product, H20, was trapped from the helium car-
rier gas. Similarly, the purified methane was con-
verted to CO 2 and H20 by combustion over hot
copper oxide at 830°C. The H20 was separated
from the CO 2 in a dry ice-cooled trap. These H20
samples were reduced to hydrogen gas over hot 43
uranium and trapped on an active charcoal at
liquid nitrogen temperatures [16]. The D/H ratios
of gases were then analyzed on a mass spectrome-
ter.
The isotopic ratios were expressed as 8 values
relative to the SMOW standard. Overall accuracies
of water samples are estimated to be _ 2%o for 8D
and _+0.2%0 for 8180. The overall reproducibility
for hydrogen and methane D/H analyses is _+ 5
and _ 3%o, respectively.
4. Results and discussion
4.1. Chemical composition of volcanic and geother-
mal gases
The analytical results (in vol.%) of the gases
sampled from volcanic and geothermal areas are
given in Table 1 and 2. Fig. 2a is a triangular
diagram showing the relative CO2, H2S + SO 2 and
N 2 contents exclusive of the major constituent
H20. In most of the areas, carbon dioxide con-
stitutes a large proportion of the gases. Hydrogen
sulfide is a predominant sulfur species except for
Za~. The Manza karabuki fumarolic gases in
Kusatsu shirane are characterized by exceptionally
high H2S contents relative to other components.
Little or no HCI is contained in these volcanic and
geothermal gases. The residual non-acidic gases of
Nasu and Issaiky~ (Azuma) have appreciable con-
centrations of hydrogen and methane, while the
gas in other volcanic areas has a high proportion
of nitrogen (Fig. 2b). In the Matsukawa geother-
mal system, hydrogen and methane concentrations
are much higher than those in other volcanic re-
gions. In most of the gas samples from these
volcanic and geothermal areas, the content of hy-
drogen is higher than that of methane.
High-temperature volcanic gases predominantly
contain SO 2, HC1 and H 2 except for H20 and
CO 2, and little CH 4 [17-19]. When these deep-
seated volcanic gases are cooled and condensed
through groundwater, sulfate and hydrogen sulfide
are formed by the disproportionation of SO 2, and
hydrogen chloride is dissolved in water [20]. Up-
ward movement of this thermal system containing
gases is accompanied by boiling and phase separa-
tion. Consequently, sulfate and chloride are dis-
44
TABLE 1
Chemical composition of volcanic gases collected 1981 at Zag, Azuma (Issaiky~), Nasu and Kusatsu shirane from June 1980 to
October 1981 (in vol.%).
Location Outlet H20 CO 2 H2S SO 2 HCI N 2 H 2 CH 4 Ar
T(°C) (xl0 -4) (xl0 -4) (xl0 -4)
Za~ 94.3 96.8 2.51 0.018 0.27 tr. 0.398 3.9 21 25
96.0 97.7 1.98 0.012 0.24 tr. 0.088 4.3 4.1 4.0
Issaiky~ 87.0 - 86.6 7.33 0.30 - 5.78 607 63.6 231
94.0 97.0 2.30 0.537 0.099 tr. 0.054 61.2 17.0 1.7
Nasu Oana 110 99.6 0.224 0.064 0.009 0.016 0.053 130 71.7 4.7
112 99.7 0.210 0.065 0.013 0.007 0.008 27.4 4.5 0.7
113 99.7 0.213 0.066 0.014 0.012 0.006 14.3 2.2 0.3
Nasu Mugen 149 99.6 0.298 0.052 0.012 0.006 0.017 26.2 15.1 3.2
169 99.8 0.156 0.037 0.002 0.011 0.005 7.0 1.1 0.6
154 99.7 0.243 0.046 0.005 0.011 0.004 15.8 2.0 0.3
Kusatsu shirane 96.0 98.2 1.35 0.196 0.009 tr. 0.439 18.9 1.2 28.1
Kusatsu
seshogawara 95.0 94.1 3.55 2.11 0.01 tr. 0.225 5.1 39.5 16.3
Kusatsu yubatake 64.0 - 49.6 1.89 0.12 - 47.5 29.5 987 8320
Manza karabuki 96.0 99.2 0.051 0.702 0.008 tr. 0.039 5.3 2.8 5.2
94.0 99.2 0.146 0.629 0.022 0.002 0.004 0.54 0.15 0.36
tr. = trace;- = not determined.
tributed in the liquid phase and H2S, CO 2 mainly
into the vapor. Hydrogen remains together with
methane in the gas phase because these gases are
hardly dissolved in the liquid phase. This liquid
phase may correspond to acid chloride-sulfate type
waters which exist in northeastern Japanese
volcanic areas such as Tamagawa, Za~ and Kusatsu Manza. Thus, a mixture of the high-temperature
gases with groundwater is indicated by little or no
HC1 and SO 2 contents, and high H2//CH4 ratio in
most gas analyses. The high water content is also
again suggestive of meteoric water contamination.
Therefore, these facts suggest that most of the
samples collected in northeastern Japan represent
TABLE 2
Results of chemical and isotopic analyses of collected gases at Matsukawa geothermal areas (in vol.% and %o, respectively)
Well No.: 1 2 2 5 6 8 8 9
Year: 1977 1977 1981 1981 1977 1977 1981 1981
H20 99.5 99.6 99.8 99.7 99.6 99.7 99.6
CO 2 0.415 0.319 0.201 0.296 0.305 0.199 0.320
H 2 S 0.043 0.051 0.031 0.036 0.055 0.065 0.069
N 2 ( × 10 -4) 49.9 25.2 22.7 37.2 47.2 20.5 43.2
H 2 ( X 10 -4) 24.1 63.6 38.5 14.4 14.0 29.7 38.6
CH 4 (X l0 -4) 43.2 10.9 10.8 15.3 35.1 5.5 22.2
DH2 -- 364 - 419 - 465 - 399 - 385 - - 406
8 DCH 4 -- 274 -- 386 -- 410 -- 351 -- 274 -- 320 -- 384
8 DH2 o -- 76.3 -- 76.8 -- 74.1 -- 78.3 -- 75.8 -- 72.8 -- 79.5
818OH2 o -- 4.8 -- 4.1 -- 7.8 -- 8.4 -- 5.0 -- 5.3 -- 8.8 99.4
0.538
0.033
25.0
47.6
20.0
- 466
- 395
- 76.6
-8.1
- = not determined.
45
(b) N2
20 o 80
®
CO 2 " ).,p~.. 80 s0 t0 20 ~s" r'°2
L \
,0.,7".'.." 80 60 40 20 N2
b the composition of gas phase equilibrated with
heated groundwater.
4.2. Hydrogen isotopic composition of volcanic hy-
drogen
The results of isotopic analyses of volcanic hy-
drogen and methane are listed in Table 2 and 3,
and shown in Fig. 3. The variation of 8D in
volcanic hydrogen is considerable; -220 to
-515%o. Volcanic hydrogen from Nasu fumarolic
gases, in which HOI is found in low concentration
at temperatures in the range from 110 to 170°C, is
depleted in deuterium compared to other Japanese
volcanic gases. On the other.hand, the 8D value of
the hydrogen at Kusatsu shirane, in which the
ratio of hydrogen to methane is large, is -2207~
Fig. 2. (a) Relative proportions of carbon dioxide, and total
sulfur and nitrogen in volcanic and geothermal gas sa~aples,
northeastern Japan. (b) Relative H2, CH 4 and N 2. • = Zao,
Q= Azuma (IssaikyS); ® = Nasu; O = Kusatsu shirane; @=
Matsukawa.
TABLE 3
Isotopic analyses of hydrogen, methane and waters collected at Za~, Issaiky~, Nasu and
1981 (in 96o) Kusatsu shirane from June 1980 to October
Location 8DHz 8DcH 4 8OH: O 818OHIo
Za~ - - - 87.0 - 13.6
- 67.3 - 9.8 (h.w.)
-370 - 181 -77.1 - 10.8
- 58.9 - 7.3 (h.w.)
Issaiky~ - - - 91.4 - 12.8
- 486 - 378 - 69.7 - 8.0
Nasu Oana -488 -400 -62.3 -7.1
- -394 -59.5 -6.8
-515 -430 -59.1 -6.8
Nasu Mugen - - 457 - 63.9 - 6.7
-31.4 - 1.3 (h.w.)
- -407 -61.4 -6.8
- - 487 - 60.5 - 6.7
Kusatsu shirane - 220 - - 75.9 - 9.3
Kusatsu seshogawara - -220 - 102 - 16.6
Manza karabuki -300 - - 101 - 14.5
- 72.7 - 8.7 (h.w.)
- - - 103 - 15.0
-73.1 -8.7 (h.w.)
h.w. = hot water; -: not determined.
46
Surtsey
Zo5
Azuma
NOsu
Kusotsu shirane
Matsukawa
IceLand
Yellow stone t
0
i
I
i
• o I
I
I
I
I
o
I
i
i
n j.J-L o IF.=.
I I I I I I
- 600 - 400 - 200 ~ D (%,)
Fig. 3. Hydrogen isotopic ratios of hydrogen and methane in
volcanic and geothermal gases. Also included are data on
hydrogen from Surtsey volcano [1] and Iceland geothermal
areas [2], and hydrogen and methane from Yellowstone hot
springs [4]. Large solid and small open circles indicate hydro-
gen and methane, respectively.
and much higher than those of other volcanic
areas except Surtsey [1]. The average 8D value of
volcanic hydrogen from northeastern Japan is
much lower than those of hydrogen from the
Surtsey volcano in Iceland. The hydrogen in hot-
water-dominated geothermal system in Iceland is
significantly depleted in D compared with that
from volcanic gases in Surtsey and the northeast-
ern Japanese volcanic areas, but the isotopic val-
ues are higher than those of the high-temperature
geothermal area at Yellowstone Park (SD about
-700%o [4]). On the other hand, the 8D value of
hydrogen from several wells in Matsukawa, which
range from -360 to -470%~, are higher than
other geothermal areas and are similar to those
from volcanic gases in northeastern Japan except
for Kusatsu shirane.
4.3. Isotopic composition of waters from volcanic
and geothermal areas
On the basis of isotopic studies, it has been
concluded that most waters associated with
volcanic and geothermal areas are of local meteoric
origin [21,22]. Kusakabe et al. [22] and Mizutani
[23] gave isotopic data on fumarolic condensates j/ -40
~-60
t-~
oo
-80
4
Hot ~k:lt er
Local Meteoric
o Water -i oo
I
-1 5 -10 ~ls 0 (,/,,) -5 0
Fig. 4. Isotopic composition of waters from volcanic and geo-
thermal areas. M.W.L = Meteoric water line (gD = 8 8n80+
10). • = Za6; Q= Azuma (Issaiky~), • = Nasu; © = Kusatsu
shirane; ~ = Matsukawa.
from several volcanoes of Japan. The results indi-
cate that Japanese volcanic condensates are de-
rived from local surface water which has been
heated and enriched in ~80 by exchange with
silicates and/or a mixture of local meteoric waters
with magmatic water recycled from hydrous crustal
rocks. The former situation best describes the
low-temperature fumarolic condensates from most
of the volcanoes in this study, the latter, the
high-temperature (> 400°C) steam condensates
such as Showashinzan [18,23] and Satsuma-
Iwojima [19].
The relationship between 8D and 180 of vari-
ous type of waters collected in this study is shown
in Fig. 4. The pattern found in this figure is simi-
lar to those in the low-temperature fumarole in
other volcanoes as above. Volcanic condensates
from low-temperature fumaroles of -100°C in
Za6 and Manza karabuki (Kusatsu shirane) are
accompanied by boiling water and liquid-vapor
separation. The temperatures of liquid-vapor sep-
aration in these fumaroles estimated from the ex-
tent of fractionation between steam and hot water
agree with the fumarolic temperature. Most steam
condensates with small D and 180 shifts are essen-
tially local meteoric in origin. It will also be seen
from Fig. 4 that the deuterium content of steam at
Matsukawa vapor-dominated geothermal systems
is approximately equal to that of the local meteoric
water. However, the enrichment in 180 of steam
can be attributed to high-temperature exchange
with silicates.
4.4. Isotopic geothermometry
Volcanic hydrogen is significantly depleted in
deuterium compared to water, due to hydrogen
isotope exchange. The 8 D values of coexisting H 2
and H20 have been used as a geothermometers,
assuming that the gases are in isotopic equilibrium
through the following reaction:
H20 + HD = HDO + H 2
The isotope distribution in a gas phase is given by
the fractionation factor a for the exchange reac-
tion:
a = (D/H)w/(D/H)H
where "W" and "H" indicate water vapor and
TABLE 4
Comparison of isotopic temperatures with temperatures mea-
sured in volcanic and geothermal areas
Location Otw-H 2 Isotope Outlet
temper- temper-
ature ature
(°c) (°c)
Za~ 1.4938 390 96
lssaiky~ 1.8097 248 94
NasuOana (1) 1.8315 241 110
(2) 1.9400 211 113
Kusatsu shirane 1.1847 714 95
Manza karabuki 1.3247 526 96
Matsukawa Year Otw_H2 Isotope Bottom
Well No. temper- temper-
ature ature
(oc) (oc)
1 1977 1.4523 415 250
2 1977 1.5890 342 240
2 1981 1.7307 270 240
5 1981 1.5336 370 220
6 1977 1.5028 382 240
8 1981 1.5497 370 260
9 1981 1.7292 270 230
a = fractionation factor for the hydrogen isotope exchange
between water vapor and hydrogen gas. 47
hydrogen, respectively. The equilibrium fractiona-
tion factor between water vapor and hydrogen is
taken from both theoretical and experimental
studies [24-26]. According to Rolston et al. [26],
the fractionation factor, a, over the range
273-473°K can be expressed as:
In a = -0.2735 + 449.2/T+ 2380/T 2 (1)
where T is the absolute temperature. This experi-
mental scale is nearly consistent with the theoreti-
cal scale by Bottinga [24]. Hydrogen isotopic ex-
change could also taken place between hydrogen
and other hydrogen compounds such as HzS and
CH 4. However, since water vapor constitutes more
than 95% of the volcanic gases in northeastern
Japan, the isotopic behavior of hydrogen would be
controlled by isotopic equilibrium with water. Tak-
ing the 8D value of meteoric water in northeastern
Japan [27] and the isotopic fractionation factors
between water vapor and hydrogen at a maximum
temperature range of 100 - 400°C [26], hydrogen
of 8D = - 360 to - 650%o in volcanic and geother-
mal areas would be estimated (Fig. 3). The 8D
value of hydrogen becomes lower when the tem-
perature of exchange reaction falls. Except for the
fumarolic gases in Kusatsu shirane, the observed
values of -364 to -515%o are compatible with
the model if the uncertainties in temperature is
taken into account. This fact suggests that volcanic
hydrogen and water are in isotopic equilibrium
with each other in a hydrothermal temperature
range (100-400°C).
Isotopic temperatures calculated for isotopic ex-
change between water vapor and hydrogen using
equation (1) are summarized in Table 4. The iso-
topic temperatures in Za~ and Manza karabuki
(Kusatsu shirane) are calculated from the deu-
terium content of local surface water because in
both areas the exchange is accompanied by
liquid-vapor separation. The isotopic temperatures
obtained range from 200 to 700°C, and are much
higher than the outlet temperature of the fumaroles
of around 100°C. The highest outlet temperature
of Mugen fumaroles in Nasu volcano of 180°C
nearly agrees with the isotopic temperature. Arna-
son [2] obtained reasonable agreement between the
measured temperatures in wells of Iceland and the
isotopic temperatures from the hydrogen-water ex-
48
change reaction. Similar results have been ob-
tained for Wairakei, New Zealand [28]. Therefore,
the discrepancy between observed and calculated
temperatures indicates that the high equilibrium
temperature at depth has been frozen-in as tem-
peratures decreased in the rising gas, or that iso-
topic equilibrium was never established. In the
Matsukawa geothermal areas, the isotopic temper-
atures from drill holes 2 and 9 are approximately
the same as the temperatures measured. However,
the isotopic temperatures from other drill holes are
higher than the measured temperatures. It is possi-
ble that the equilibrium isotopic ratio of hydrogen
changes due to cooling during upward flow. Kiyosu
[29] reported that the isotopic temperatures ob-
tained from 8348 of hydrogen sulfide and anhydrite
in the Matsukawa area ranged from 250 to 350°C.
These isotopic temperatures are similar to the hy-
drogen isotopic temperatures. This fact suggests
that in this area the H20, H 2 and H2S have
chemically and isotopically equilibrated with the
deeper mineral components at a temperature of at
least 250°C.
4.5. D / H ratios of volcanic methanes
The 8D values of volcanic methane are more
variable than those of hydrogen; - 180 to - 487%o,
as shown in Fig. 3. Gunter and Musgrave [4] re-
ported 8D values of -225 to -292%o for hot
spring gases of yellowstone Park. On the other
hand, the range of geothermal methanes from
Matsukawa is -247 to -410%o. The variation in
D/H ratios is remarkable and comprises a range
of 300%o. The increase or decrease of deuterium in
volcanic methane could be due to the fractionation
of hydrogen isotopes and/or the mixing of various
methanes (e.g., biogenic or thermogenic CH4).
If methane is in isotopic equilibrium with water
vapor, the H20-CH 4 hydrogen exchange reaction
is responsible for the D/H variation of volcanic
methane. At isotopic equilibrium between CH 4
and H20, for example at a temperature range of
100-400°C, the methanes should be 80-100%o de-
pleted in deuterium as compared to cogenetic water
vapor [30]. Most waters associated with volcanic
gases lie within the range - 50 to - 90%o. Methanes
in equilibrium with these waters should range in -40
- 6O
-80
-I O0 @
I I
-400 -300 /I .u
/fl,
°ti!'
,"I! ° I
-200
Fig. 5. Relationship between 8D values of methane and water.
• = Za~, Q= Azuma (Issaiky~); • = Nasu; © = Kusatsu
shirane; O = Matsukawa. Solid and dashed lines: equilibrium
isotherms of liquid water and water vapor [30].
their 8D values from -150 to -200%o. On the
other hand, the 8D values of methanes based on
the theoretical scale of Bottinga [24] range from
-96 to -155%o. However, the range of the pre-
sented volcanic gases is - 300 to - - 450%0 except
for Za~ and Kusatsu shirane, as shown in Fig. 5.
The data presented here thus show that hydrogen
isotopic equilibrium has not been achieved in the
methanes from volcanic areas.
Fig. 6 shows the relationship between the hy-
drogen isotopic composition of hydrogen and
methane from volcanic gases. The equilibrium iso-
therms based on the experimental scale of Horibe
and Craig [31] are given in this figure. The arrow
indicates the direction of the change in 8D values
in hydrogen and methane after equilibration with
the associated H20 with 8D values of -50 to
-90%o. Most data plot within the field deviated
from this arrow. This suggests that volcanic
methane has not isotopically equilibrated with
water vapor and hydrogen. On the other hand, the
higher the 8D value of volcanic hydrogen, the
/ /
I DeCrt, a$ing //H20(V ) , , / lncr~tsmg .~
-2ooL ..",
/ / / J /¢'-~/ / I I i I //'71."
4 °°t ,' ," ," ,7,'I" ~=~ I I I" , "/',7 I I I C~ / / ' (250"C) ~/// ~_'e / ~o I I i i l ® /i /" o/(3oo-4oo'o
-400 I ii ii I 111¢~7e/~
I II il i I /(~3~1
V lilt/Ill/lit ,
-600 -500 -400 -300
DR2 (%*)
Fig. 6. Plot of 8D values for hydrogen versus methane. O =
Yellowstone [4]; • = ZaB; Q= Azuma (Issaiky~); Q = Nasu;
O = Matsukawa. Bracket represents the isotopic temperature
based on the observed HzO-H 2 hydrogen isotope fractiona-
tions. Dashed line: equilibrium isotherms [31]. H20(1 ) = liquid
water. H20(v ) = water vapor.
higher the deuterium content of methane is. Fur-
thermore, the D concentration in the methane
increases with increasing isotopic temperature
calculated from H20-H 2 hydrogen isotope frac-
tionations as shown in Fig. 6. Thus, D/H isotope
analyses of these methanes reveal that their hydro-
gen isotopic composition are dependent on the 8D
value of hydrogen and isotopic temperature.
Variation of 8D values in volcanic methanes
may be caused by admixture of natural gases of
different origin. According to Schoell [32], the 8D
values of biogenic methane from natural gases of
worldwide occurrences, range from -180 to
-2807o0. It should be noted that the D/H varia-
tions in volcanic methanes of northeastern Japan
are not solely due to the mixture of biogenic
methanes, since most of the 8D values for volcanic
methanes are more negative than -300~. Al-
though the hydrogen isotopic composition of ther-
mogenic methane ranges from a 8D value of - 150
to -260700 [32], their D/H variations are qualita-
tively comparable to the variations which were
observed in laboratory pyrolysis experiments [16].
At 400 ° and 500°C there are hydrogen isotope
fractionations during methane production of 170
and 140%o, respectively, with deuterium depletion
in the methane caused by the by pyrolysis of 49
hydrocarbons. Although there is an uncertainty, a
temperature effect of 30%o per 100°C in the pre-
dictable direction is obtained, If this fact is cor-
rect, then hydrogen isotopic fractionation during
methane production below 300°C would be larger
than 200%0. Assuming that hydrocarbons of a
mean isotopic composition such as that of crude
oils [33] and kerogens [34] were the parent materials
to all these gases, the deuterium concentrations of
produced methane range from - 350 to - 4007~ at
300°C, -230 to -320%o at 400°C and -180 to
- 3007o0 at 500°C. Thus, the 8D values in thermo-
genic methane increase with increasing tempera-
ture. This variation pattern is similar to that of the
values observed. Fig. 7 shows a plot of 8D value
for methane vs. the CH4/H: ratio. The deuterium
concentrations are related to the CH4/H 2 ratio
and would also increase with increasing isotopic
temperature as above. The production of methane
by pyrolysis of hydrocarbon occurs rapidly as
temperature increases [16]. Therefore, D/H varia-
tion of volcanic methane may be due to the kinetic
effect during the production of thermogenic
methane.
The 8D values of methane may also be affected
by the kinetic effect during inorganic production,
assuming that the chemical reaction proceeds as
-200
~ -300
t"~
t,,o
- 400 ee '1
I
e/
I
0.1 1
CHdH 2 I
10
Fig. 7. Plot of 8D values for methane versus CH4/H: ratios.
• = Za~; ®= Azuma (IssaikyB); (D = Nasu; O = Kusatsu
shirane; O = Matsukawa.
50
follows:
CO 2 + 4 H 2-- CH 4+ 2 H20
This chemical reaction is important for the forma-
tion of methane [3,35] and has been suggested in
early carbon isotope studies [36,37]. Schoell [32]
and Coleman et al. [38] reported D/H fractiona-
tion to occur when methane is produced and/or
oxidized by microbial processes. Isotopic
fractionation in the inorganic formation or oxida-
tion of methane may also be controlled by kinet-
ics. This reaction shifts to the right with decreasing
temperatures. Since hydrogen readily exchanges
isotopically with the associated water vapor and
methane, the 8D values of hydrogen and methane
decrease as the CHa/H 2 ratio increases in re-
sponse to a decrease in temperature. However,
lower CH4/H 2 ratio should result in even more
negative 8D values of hydrogen and methane pre-
sented here as shown in Figs. 6 and 7. Therefore,
the volcanic methanes may not be formed by the
reaction of carbon dioxide and hydrogen. Thus,
this formation process is obviously of minor im-
portance for the origin of volcanic methane.
Gunter and Musgrave [4] suggested that the
methane of hydrothermal gases from Yellowstone
Park is formed by the thermal decomposition of
organic matter, not by the reaction of hydrogen
with carbon dioxide. Variations in 8D values of
volcanic methane in northeastern Japan may be
caused by a mixture of thermogenic methane. Fur-
ther discussion may become possible when a sys-
tematic carbon isotope study of the volcanic
methane is completed.
5. Conclusions
The hydrogen isotopic composition of hydrogen
from low-temperature fumarolic gases (- 100°C)
in northeastern Japan is lighter than that of high-
temperature (> 1000°C) fumarolic gases such as
those from the Surtsey volcano in Iceland, and
havier than hydrogen of hot spring gases in Yel-
lowstone Park. The D/H ratio of hydrogen from
geothermal gases in the Matsukawa vapor- dominated system is higher than that of the Ice-
land hot-water-dominated geothermal areas and
similar to that of fumarolic gases in Nasu volcano.
The hydrogen isotope data suggest that the hydro-
gen of volcanic gases in northeastern Japan may
be isotopically equilibrated with water vapor at a
temperature of 200-400°C at depth.
In volcanic regions other than Nasu, the iso-
topic temperature, evaluated from the isotopic
fractionation factor between water vapor and hy-
drogen, is much higher than the outlet tempera-
tures of the fumaroles. The discordance between
observed and isotopic temperatures indicates the
freezing-in of a high-temperature equilibrium at
depth as temperatures decrease in the rising gas, or
the lack of isotope exchange between water vapor
and hydrogen at low temperatures. In the Nasu
volcano, the isotopic temperatures are similar to
the fumarolic temperatures, suggesting that the
water vapor equilibrates isotopically with hydro-
gen near the surface. The isotopic temperatures
from Matsukawa geothermal area are higher by
30 - 105°C than the bottom temperatures in drill
holes.
The hydrogen isotope ratio of volcanic methane
is more variable than that of hydrogen. The D/H
isotope data also suggest that most volcanic
methane has isotopically disequilibrated with water
vapor and hydrogen. Variation of 8D values in
these methanes may be due to a mixture of ther-
mogenic methane. Future investigations of carbon
and hydrogen isotopes will be most useful in order
to clarify the origin of volcanic methane.
Acknowledgements
The author wishes to thank N. Nakai and M.
Kurahashi (Nagoya University) and Y. Yoshida
(Japan Metals-Chemicals Co.) for their help in the
collection and analysis of the gas samples, and H.
Craig (University of California), who reviewed the
manuscript, for helpful comments.
Part of the cost of this study was defrayed by a
Grant in Aid for Scientic Research, No. 554168,
from the Ministry of Education.
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|
(JAPAN) Kiyosu 1982 - d2H-H2 and d2H-CH4 from volcanic areas in northeastern Japan.txt
|
The Island Arc
(2003)
12,
190–206
Blackwell Science, LtdOxford, UK
IARThe Island Arc1038-48712003 Blackwell Science Asia
122June 2003
390
Paleozoic ophiolites and blueschists
A. Ishiwatari and T . Tsujimori
10.1046/j.1038-4871.2003.00390.x
Original Article190206BEES SGML
*Correspondence.
Received 13 May 2002; accepted for publication 8 January 2003.
© 2003 Blackwell Publishing Asia Pty Ltd.
Thematic Article
Paleozoic ophiolites and blueschists in Japan and Russian Primorye in the
tectonic framework of East Asia: A synthesis
A
KIRA
I
SHIWATARI
1,
*
AND
T
ATSUKI
T
SUJIMORI
2
1
Department of Earth Sciences, Faculty of Science, Kanazawa University, Kanazawa 920-1192, Japan (email:
geoishw@kenroku.kanazawa-u.ac.jp) and
2
Research Institute of Natural Sciences, Okayama University of Science,
Okayama 700-0005, Japan
Abstract
Ophiolites and high-pressure (HP) metamorphic rocks are studied to test
continuation of Paleozoic and early Mesozoic geological units from Japan to Primorye overthe Japan Sea. The early Paleozoic ophiolites are present on both sides, and the latePaleozoic ophiolite of south-western Japan may also have its counterpart in Primorye. The
Shaiginskiy HP schist and the associated Avdakimov gneiss in Primorye, both tectonicallyunderlying the early Paleozoic ophiolitic complex, yield a 250-Ma phengite and hornblendeK–Ar age, which is intermediate between those of the Renge (280–330 Ma) and Suo (170–220 Ma) blueschists in south-western Japan. This age also coincides with that of the coesite-bearing eclogites in the Sulu–Dabie suture in China and several medium-pressuremetamorphic rocks in East Asia. On the basis of these results and other geological data, theauthors propose the ‘Y aeyama promontory’ model for an eastward extension of the Sulu–Dabie suture. The collision suture warps southward into the Y ellow Sea and detours aroundKorea, turns to the north at Ishigaki Island in the Y aeyama Archipelago of Ryukyu, where
it changes into a subduction zone and further continues toward south-western Japan andPrimorye. Most ophiolites from this area represent crust–mantle fragments of an islandarc–back-arc basin system, and the repeated formation of ophiolite–blueschist associationsmay be due to the repetition of the Mariana-type non-accreting subduction and Nankai-type accreting subduction.
Key words:
Japan Sea, Khanka terrane, Korea, Sikhote Alin, Sulu–Dabie suture, Y aeyama
promontory.
INTRODUCTION
More than a half century ago, long before theestablishment of plate tectonics, Kobayashi (1951)proposed a rifting–drifting hypothesis for the ori-gin of the Japan Sea. From a geological point ofview, the rifting–drifting theory requires theoccurrence of equivalent pre-Tertiary geologicalunits on both sides of the Japan Sea. For example,the Appalachian and Caledonian belts on bothsides of the Atlantic Ocean, respectively , representseparated fragments of a single early Paleozoicorogenic belt, and include Early Ordovician ophi-
olites of identical age (Dunning & Pedersen 1988).
Recent paleomagnetic results indicate a fast
drifting of Japan at
ca
15 Ma in the manner of the
opening of a pair of hinged doors (Otofuji &Matsuda 1984; Otofuji
et al
. 1985). Some basalt
samples drilled from the Japan Sea floor have
40
Ar-
39
Ar ages of 15–25 Ma (Kaneoka
et al
. 1992). These
data coupled with other geophysical and geochem-ical data result in various models for the Mioceneback-arc opening process (Nohda
et al
. 1988;
Tamaki & Honza 1991). However, evidence and
consideration for an original geological continuitybetween Japan and Russian Primorye are notconclusive.
The Japanese Islands are mainly composed of
accretionary complexes of Paleozoic, Mesozoic and
Paleozoic ophiolites and blueschists
191
Cenozoic ages. Every accretionary complex is
characterized by the ‘oceanic plate stratigraphy’,which is composed of fragments of oceanic crustand seamounts, chert and/or pelagic limestone, sil-iceous shale, sandstone and conglomerate (or olis-tostrome) in younging order (Isozaki 1996).Kojima (1989) pointed out that the Jurassic accre-
tionary complexes showing the same age–lithologyrelationship are found on both sides of the JapanSea, namely in south-western Japan (Mino-Tambabelt) and in the Sikhote Alin terrane in Primorye(Samarka zone) and the adjacent Chinese terri-tory (Nadanhada zone). However, detailed com-parison of Paleozoic ophiolites and blueschists onboth sides of the Japan Sea has not previouslybeen attempted.
The present paper reports petrologic and geo-
chronologic similarity of the Paleozoic igneous andmetamorphic rocks on both sides of the Japan Seaon the basis of our recent cooperative works withRussian geologists, and discusses configuration of
the late Paleozoic–early Mesozoic collision sutureand geotectonic significance of the multiple ophio-lite–blueschist assemblages in East Asia. Numericage data and detailed petrology of the dated sam-ples will appear elsewhere.
PALEOZOIC OPHIOLITES AND BLUESCHISTSIN JAPAN
SOUTH-WESTERN JAPAN
The Japanese Islands bear ophiolitic complexes ofvarious ages ranging from early Paleozoic to Cen-ozoic, forming a Phanerozoic multiple ophiolitebelt (Ishiwatari 1991, 1994). Paleozoic ophiolites(Fig. 1) occupy a higher structural position in thepiling nappes of the accretionary complexes. Insouth-western Japan (Fig. 2) the Oeyama ophioliteof Cambro-Ordovician age occupies the higheststructural position, and tectonically overlies thelate Paleozoic Renge blueschist, the Y akuno ophi-olite, and the Permian Akiyoshi (and Ultra-Tamba)accretionary complexes. These tectonic units arebounded by thrust faults (Ishiwatari
et al
. 1999).
The Oeyama ophiolite is mainly composed of
residual peridotite with podiform chromite depos-its and minor gabbroic rocks; basaltic volcanicrocks are completely absent. The peridotite is lher-zolitic in the eastern part (Oeyama body , spinel Cr#(Cr/(Al
+
Cr))
=
0.3; Kurokawa 1985), but is
harzburgitic in the western part (Tari-Misaka body ,spinel Cr#
=
0.5; Arai 1980). The peridotite includes
metagabbro and amphibolite bodies, which have
Fig. 1
Distribution of Paleozoic ophiolites
and Paleozoic–early Mesozoic blueschists inJapan.
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192
A. Ishiwatari and T. Tsujimori
hornblende K–Ar ages of 444–464 Ma (Nishimura
& Shibata 1989). The kyanite- and staurolite-bearing HP metagabbro (troctolitic cumulate ori-gin) of higher metamorphic pressure in the Oeyamaperidotite body also has a similar hornblende K–Arage of 403–443 Ma (Tsujimori 1999; Tsujimori
et al
.
2000c; Tsujimori & Ishiwatari 2002).
The Y akuno ophiolite consists of a relatively
complete succession composed of harzburgite tec-tonite (spinel Cr#
=
0.6–0.8), dunite–wehrlite–
clinopyroxenite cumulate, metagabbro, amphibo-lite and metabasalt with abundant black shale(Ishiwatari 1985a). Middle–Late Permian radiolar-ian fossils were identified from the black shaleintercalated with basalt lavas (Kurimoto & Makim-oto 1990). Hornblende K–Ar dates of metagabbrorange from 241
±
12 to 278
±
10 Ma (Shibata
et al
.
1977), and conventional zircon U–Pb ages of theplagiogranite are 282
±
2 and 285
±
2Ma (Herzig
et al
. 1997), indicating an Early Permian igneous
age and slightly later metamorphism for this ophi-olite. In contrast, Sano (1992) reported a 421
±
54-
Ma Nd–Sm whole-rock isochron age (8 points) forthe metagabbro and plagiogranite, and a 311
±
65-
Ma Nd–Sm rock–clinopyroxene isochron age forthe metabasalt, suggesting a polygenetic naturefor the Y akuno ophiolite. However, coincidence ofthe Nd–Sm age of the Y akuno metagabbro and theK–Ar age of the Oeyama metagabbro may notindicate their cogenetic relationship. They shouldbe compared with the age data for the samemethod. Herzig
et al
. (1997) point out that the iso-
chron from Sano (1992) might be a mixing line.
The basaltic and gabbroic rocks of the Y akuno
ophiolite show transitional mid-ocean ridge basalt(T-MORB) chemistry in the eastern part and arc–tholeiite chemistry in the western part (Ishiwatari
et al
. 1990a). The metamorphic grade increases
from the prehnite–pumpellyite facies in the upperbasalt section to the granulite facies in the lowergabbro and ultramafic section. The granulite-facies metacumulate at the Moho is a gneissose
Fig. 2
Geological units in the main part of the Inner Zone of south-western Japan. Kyoto–Tottori cross-section is shown in the inset. ISTL, Itoigawa–
Shizuoka Tectonic Line; MTL, Median Tectonic Line; MB, metamorphic belt; AC, accretionary complex. Unit boundaries are mainly a fter Ishiwatari (1994),
Tsujimori (1998), and Geological Survey of Japan (1992).
14401738, 2003, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.2003.00390.x by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Paleozoic ophiolites and blueschists
193
metagabbro composed of aluminous two
pyroxenes, plagioclase, and aluminous spinel(Ishiwatari 1985b). Such an aluminous-spinel met-agabbro is rare among ophiolitic complexes, andhas so far been reported only from Bikin, Pri-morye (Vysotskiy 1994) and Tonsina, Alaska(DeBari & Coleman 1989), although analogousrocks are known from meta-igneous complexes oflower continental crust (Rivalenti
et al
. 1981) and
lower crustal xenoliths in island arc (Francis 1976)and oceanic plateau (Grégoire
et al
. 1998).
The Renge blueschist occurs either as thin tec-
tonic slices or blocks in serpentinite mélangesunderlying the Oeyama ophiolite (Tsujimori 1998),and has a phengite K–Ar age of 320 Ma (Tsujimori& Itaya 1999; Fig. 3). The metamorphic assemblageof the mafic schist ranges from lawsonite blueschistto epidote blueschist and further into eclogitic rocksin the Omi area (Tsujimori
et al
. 2000a,b; Tsujimori
2002). The Renge metamorphic belt also includessome relatively low-pressure, high-temperaturemetamorphic rocks including oligoclase–biotiteschist and amphibolite. The Joetsu metamorphicbelt also bears typical blueschist and relativelyhigh-temperature pelitic schists with phengite K–Ar ages of 308 and 289 Ma (Y okoyama 1992), and isthought to be an eastern extension of the Renge belt(Fig. 1).
The Suo metamorphic belt also consists of high-
pressure (HP) metamorphic rocks includingpumpellyite–actinolite schist and epidote blue-schist (mostly winchite schist; Nishimura 1998).Some pelitic rocks bear lawsonite (Hayasaka 1987;Watanabe
et al
. 1987, 1989) but lawsonite–
glaucophane and pumpellyite–glaucophane assem-blages are absent. Phengite K–Ar ages are220
±
7Ma in the Nishiki area and 170–190 Ma in the
eastern areas (Nishimura 1998; Fig. 3). The schistswith the same K–Ar age are also known from theKurosegawa belt of the Outer Zone of south-west-
ern Japan and Ishigaki Island of Ryukyu (Fig. 3).
NORTH-EASTERN JAPAN
Early Paleozoic Miyamori ophiolite forms thebasement of a nearly complete Paleozoic–Mesozoicsedimentary sequence of the South Kitakami beltranging from Ordovician to Jurassic in age(Tazawa 1988, 2000). The Miyamori ophiolite con-sists mostly of wehrlitic cumulate and depletedmantle harzburgite (spinel Cr#
=
0.4–0.8) with
hornblende and phlogopite, which is interpreted tohave been upper mantle of an island arc (Ozawa1988). Minor lherzolite patches (spinel Cr#
=
0.1–0.4) of kilometric sizes are included in the depleted
harzburgite, and are interpreted to have been arelict source mantle that originated in a back-arcbasin (Ozawa 1988). Hornblende K–Ar ages of fourgabbroic rocks cutting peridotite of this ophiolite
Fig. 3
Isotopic ages of ophiolitic and metamorphic rocks in south-
western Japan, Russian Primorye and related areas. (a) K–Ar ages ofhigh-pressure schists in Primorye and Japan. Data sources: Kovalenkoand Khanchuk (1991); Tsujimori and Itaya (1999); Nishimura (1998) andthe references therein. Data for Primorye are mostly based on our unpub-lished results.
14401738, 2003, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.2003.00390.x by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
194
A. Ishiwatari and T. Tsujimori
range from 445 to 485 Ma, although other horn-
blendite (421 Ma) and amphibolite (369 Ma) yieldyounger ages (Ozawa
et al
. 1988). These data indi-
cate an Ordovician or earlier age for the Miyamoriophiolite, the same age as the Oeyama ophiolite insouth-western Japan.
The Motai blueschist belt (Maekawa 1988) in the
South Kitakami belt also bears pyroxene amphib-olite blocks with hornblende K–Ar ages of 479
±
24
and 524
±
26 Ma (Kanisawa
et al
. 1992). The K–Ar
age of a calk-alkaline tonalite dike cutting the meta-morphic rocks is 457 Ma (Sasada
et al
. 1992). A
monazite chemical Th-U-total Pb isochron method(CHIME, or electron microprobe method) age of430
±
10 Ma is also reported from the paragneiss in
the 350-Ma granitic complex (Suzuki & Adachi1991). These data indicate that ophiolites and meta-morphic and granitic rocks comprised an early
Paleozoic active continental margin or mature
island arc. The 300-Ma muscovite K–Ar age(Kawano & Ueda 1965) and 225
±
11- and 239
±
12-
Ma hornblende K–Ar ages (Kanisawa
et al
. 1992) of
garnet–epidote amphibolites from the Y amagamiarea suggest that the blueschist metamorphism inthe southern part of the Motai belt is contemporarywith the Renge and Suo blueschists of south-western Japan. However, late Paleozoic ophiolite(Y akuno) and accretionary complexes (Akiyoshiand Ultra-Tamba) are absent in north-easternJapan. The early Paleozoic ophiolitic–granitic base-ment and the sedimentary cover of Silurian–Jurassic ages in the South Kitakami belt thrustover the Jurassic accretionary complex of theNorth Kitakami belt (Tazawa 1988, 2000).
Fig. 3
(
Continued
)
(b) K–Ar ages for ophiolitic complexes in Primorye and Japan. Data sources: Khanchuk
et al
. (1996); Kovalenko and Khanchuk
(1991); Nishimura and Shibata (1989); Shibata
et al
. (1977); Sano (1992); Ozawa
et al
. (1988). K–Ar data for Primorye are mostly based on our unpublished
results. (c) Muscovite–phengite K–Ar ages of schists in the Kurosegawa belt, Outer Zone of south-western Japan (compilation in Tsujimori
et al
. 2000c).
(d) Nd–-Sm ages and zircon U–Pb ages of the ultrahigh-pressure metamorphic rocks in the Sulu and Dabie areas, China (after Ames
et al
. 1993; Li
et al
.
1993). ( ), hornblende K–Ar age; ( ) muscovite–phengite K–Ar age; ( ), muscovite–phengite Ar–Ar age; ( ), zircon U–Pb age; ( ), Nd–S m min-
eral isochron; ( ), Nd–Sm rock isochron.
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Paleozoic ophiolites and blueschists
195
The radiometric ages of pre-Cretaceous ophi-
olitic and metamorphic rocks in south-westernJapan are summarized in Fig. 3. The ophiolite agesare centered on two peaks at approximately 450 Ma(Oeyama ophiolite) and 280 Ma (Y akuno ophiolite).North-eastern Japan bears the 450-Ma ophiolitesbut lacks the 280-Ma ophiolites. The HP metamor-phic rocks are concentrated around two otherpeaks at approximately 300 Ma (Renge blueschist)and 200 Ma (Suo blueschist), although minor olderHP metamorphic rocks of 400–450 Ma are alsopresent in association with the Oeyama ophioliteand the Kurosegawa mélange in Shikoku Island.
PALEOZOIC OPHIOLITES AND BLUESCHISTS IN PRIMORYE, RUSSIA
The Primorye territory of Russia is geologicallydivided into two parts: the Khanka terrane and theSikhote-Alin terrane (Fig. 4). The Khanka terranemay be a part of a larger continental block includ-ing the Bureya and Jiamusi blocks to the north
(Khanchuk 2001), composed of Precambrian conti-nental basement covered by thick Cambrian cal-careous sediments and post-Silurian continentalsediments. The Sikhote-Alin terrane mainly con-sists of Mesozoic accretionary complexes of green-stone, chert, limestone, shale and sandstone, whichare intruded by Cretaceous granites and coveredby Cretaceous–Tertiary volcanics.
KHANKA OPHIOLITE
The Khanka terrane bears some ophiolitic bodies inthe west-north-west-trending Spassk zone, which isalmost perpendicular to the general trend of theSikhote-Alin terrane. Shcheka
et al
. (2001) describe
an ophiolitic sequence of serpentinite (harzburgite),pyroxenite, gabbro and basalt that are emplacedonto the Early Cambrian fossiliferous limestone–shale formation and covered by Middle Cambrianconglomerate including abundant detrital chro-mian spinel. The chromian spinel grains from ser-
Fig. 4
Geological units of Russian Primorye
(simplified after Khanchuk
et al
. 1996; with addi-
tion of Shaiginskiy blueschist).
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196
A. Ishiwatari and T. Tsujimori
pentinite and conglomerate are high in Cr# (0.50–
0.75) and very low in TiO
2
(
<
0.2 wt%) as those in
other ophiolites. Although the area is free fromregional metamorphism, ferrichromit rims aredeveloped in most chromian spinel grains; some areunusually rich in MnO (up to 19 wt%), suggestingocean-floor hydrothermal alteration prior to theemplacement. Shcheka
et al
. (2001) conclude that
the Khanka ophiolite may represent oceanic litho-sphere formed in a small rift zone and immediatelyemplaced onto the adjacent passive continentalmargin prior to the circum-Pacific orogeny .
SIKHOTE ALIN OPHIOLITES
Major ophiolitic complexes of the Sikhote Alin ter-rane are aligned in a north-north-east directionparallel to the eastern margin of the Khanka ter-rane. The trend also parallels the Central SikhoteAlin Fault. The ophiolitic complexes are distrib-uted in the Sergeevka, Kalinovka and Bikin areasfrom south to north.
The Sergeevka metagabbro body forms a north-
east-trending massif of 30
¥
130 km, the largest
mafic body in Primorye. Gneissose hornblendemetagabbro occupies more than 80% of the area inassociation with some granitic and troctolitic intru-sions, as well as various metamorphic rocks suchas gneiss, amphibolite and marble. Conventionalzircon U–Pb ages of 528
±
3Ma are reported for
gneissose metagabbro, 504
±
3Ma for gneissose
diorite, and 493
±
12 Ma for granite (Khanchuk
et al
. 1996). Mishkin
et al
. (1970) reported a mus-
covite K–Ar age of 529 Ma for another granite anda hornblende K–Ar age of 622 Ma for garnetamphibolite, but Tsujimori (unpubl. data) obtainedeight hornblende K–Ar ages of metagabbroswithin a narrow range between 430 and 470 Ma(Fig. 3). These hornblende K–Ar ages are similarto those of the Oeyama ophiolite in south-westernJapan and the Miyamori ophiolite in north-easternJapan (Fig. 3). Khanchuk
et al
. (1996) consider
that this body forms a part of the continental mar-gin of the Khanka block, hence it is not an ophio-lite. However, the dominantly mafic nature of thisbody , its occurrence as a nappe overlying youngerblueschist and accretionary complex, and its posi-tion located on the same line as the other Sikhote-Alin ophiolites (Fig. 4) suggest that the Sergeevkamassif is a dismembered ophiolite body .
The Kalinovka ophiolite group is composed of
three north-east-trending ophiolitic bodies ofapproximately 5
¥
40 km having an
en echelon
arrangement. These bodies are composed of dun-ite, troctolite, wehrlite, clinopyroxenite, olivine
gabbro, hornblende gabbro, plagiogranite, pillowbasalt and minor amphibolite and granite. Thechert and limestone associated with the pillow lavabear conodont fossils of Late Devonian–EarlyPermian ages (Vysotskiy 1994). The K–Ar age of
410
±
9Ma is the only reported age determined for
very K-poor hornblende (Kemkin & Khanchuk1994). Our preliminary hornblende K–Ar datingfor the metagabbro at Medvezhy Kut nearBreyevka yields 230 Ma, which is younger thanthat of the Y akuno ophiolite (Fig. 3). Vysotskiy(1994) describes an olivine–plagioclase reactionrelationship to form aluminous spinel–pyroxenesymplectite in troctolite, and Khanchuk andPanchenko (1991) report garnet metagabbro. The
tectonic superposition of the Kalinovka ophioliteover the Jurassic accretionary complex of theSamarka zone with an intervening older accretion-ary complex called the Udeka zone resembles ananalogous relationship in south-western Japan,where Y akuno ophiolite thrust over the JurassicTamba zone with the Permian Ultra-Tamba zone in
between (Kojima
et al
. 2000).
The Bikin ophiolite group is composed of three,
north–south-trending ophiolitic bodies of 1
¥
2–
3km in size; namely the Oronsky , Zalominsky and
Soldinsky bodies (Vysotskiy 1994; Vysotskiy
et al
.
1995). Dunite, harzburgite, wehrlite, orthopyrox-enite, and aluminous spinel-bearing noritic gabbroare associated with Late Permian basaltic pillowlava and siliceous volcanic rocks as well as a ser-pentinite mélange. The aluminous spinel-bearing,olivine-free gabbro with very aluminous pyroxenes(Al
2
O
3
>
8wt%) at the Moho is evidence for unusu-
ally thick oceanic crust (Ishiwatari 1985a), and itsoccurrences on both sides of the Japan Sea (in theYakuno and Bikin ophiolites) suggest original con-
tiguity of the Paleozoic ophiolite belt.
SIKHOTE ALIN BLUESCHIST
The Shaiginsky blueschist occurs as windows andthin thrust sheets beneath the Sergeevka ophioliticbody . The epidote blueschist bears crossite andbarroisite. Pelitic rocks are of a higher grade thanthe garnet zone, and some samples bear oligoclase(An
18
), although biotite is completely absent. Gar-
net preserves progressive normal zoning withdecreasing Mn toward the rim; some showsreverse zoning at the rim. Piemontite-bearing sil-iceous schist is also present near Partisansk. TheShaiginsky schists yield phengite K–Ar ages of230–250 Ma (Fig. 3). These age data lie between
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Paleozoic ophiolites and blueschists
197
those of the two HP metamorphic belts in south-
western Japan (i.e. the Renge (280–330 Ma) andSuo belts (170–220 Ma)). The Shaiginsky blueschistis associated with the ‘ Avdakimov gneiss’ composedmainly of hornblende gneiss with marble andcoarse-grained garnet amphibolite. Although aPrecambrian Rb–Sr mineral whole-rock isochronage was reported from this gneiss complex, ourhornblende K–Ar ages are centered at 250 Ma(Fig. 3), the same age as the Shaiginsky blueschistcomplex. Kovalenko and Khanchuk (1991) reporteda 255-Ma and 290-Ma phengite K–Ar age for peliticschists of the Shaiginsky complex; our K–Ar dataare centered at 250 Ma (Fig. 3). These ages areintermediate between those of the Renge (280–330 Ma) and Suo (170–220 Ma) blueschists of south-western Japan. Nevertheless, the K–Ar age of theSuo metamorphic rocks varies significantly fromarea to area (Fig. 3); some metamorphic rocks inwestern Kyushu and in the Kurosegawa Belt havethe same age as the Shaiginskiy blueschist.
GEOLOGICAL CONTINUATION FROM JAPAN TO PRIMORYE
Even if we accept the rifting–drifting hypothesis,the Pre-Japan Sea configuration of the JapaneseIslands is not easy to restore. Some authors
assume that south-western Japan was locateddirectly to the south of Primorye and on the east ofthe Korean Peninsula (with Y amato Bank inbetween) as shown in Fig. 5 (Kojima 1989;Khanchuk 2001). In contrast, the South KitakamiBelt and associated accretionary complexes ofnorth-eastern Japan should already have beenplaced alongside Primorye or between Primoryeand south-western Japan in the Early Tertiaryprior to the opening of the Japan Sea. However,paleontologic data of Paleozoic and Mesozoic for-mations in Japan indicate that the South KitakamiBelt and the Kurosegawa Belt (the Outer Zone ofsouth-western Japan) were placed in the Chinesecontinental margin to the south of south-westernJapan in Cretaceous and earlier time (Otoh &Sasaki 1998; Tazawa 2000), and displaced north-ward through fast and extensive left-lateralstrike–slip movement.
Arakawa
et al
. (2000) mention the possibility
that the Hida belt does not belong to the Sino-Korean block, but has evolved as a part of the East-
Central Asian Orogenic Belt, which is a wide accre-tionary belt extending from Primorye to CentralAsia via north-eastern China and Mongolia along
Fig. 5
(a) Continuation of geological units from south-western Japan to Russian Primorye before opening of the Japan Sea. Position of south-western
Japan follows that of Kojima (1989). (b) Major geological events in Japan and Primorye.
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198
A. Ishiwatari and T. Tsujimori
the northern margin of the Sino-Korean block.
Khanchuk (2001) suggests that the Laoelin–Grode-kovo belt and Cheongjin belt in the Russia–China–North Korea border area are possible extension ofthe ophiolite belts in south-western Japan.
PALEOZOIC-EARLY MESOZOIC BLUESCHISTS IN JAPAN AND PRIMORYE: RELATION TO THE COLLISION BELT IN CHINA
As noted in the previous section, some late Paleo-zoic–early Mesozoic blueschist facies rocks of
ca
250 (
±
100) Ma occur in southern Primorye (Shai-
ginsky blueschist: Kovalenko & Khanchuk 1991and the present study), northern Honshu (Motaibelt: Maekawa 1988; Kanisawa
et al
. 1992), west-
ern Honshu and Kyushu (Renge and Suo belts:Nishimura 1998; Tsujimori & Itaya 1999), andIshigaki Island of Ryukyu (Tomuru Formation:Nishimura
et al
. 1983; Faure
et al
. 1988). They are
mostly epidote blueschist in Primorye, Motai, Suoand Ishigaki, but typical lawsonite blueschist andeclogitic rocks occur in the Renge belt (Tsujimori1998; Tsujimori
et al
. 2000a,b). It should be noted
that HP metamorphism in Japan took place notonly in late Paleozoic–Early Mesozoic time but alsoin early Paleozoic time (Kurosegawa: Maruyama &Ueda 1974; Oeyama: Tsujimori 1999; Tsujimori
et al . 2000c) and in Cretaceous time (Sambagawa,
Nagasaki, and Kamuikotan). Thus, subduction-zone metamorphism repeatedly took place inJapan; a subduction zone also existed in Japan atca250 Ma, when collision of the Sino-Korean and
Yangtze cratons took place (Fig. 3).
DOES THE CHINESE COLLISION SUTURE GO TO KOREA?
The Chinese Dabie-Sulu ultrahigh-pressure
(UHP) metamorphic belt is believed to be a colli-sional suture between the Sino-Korean andYangtze blocks, which amalgamated during the
200–250-Ma period, on the basis of the Nd–Sm andU–Pb ages of the UHP metamorphic rocks (Ameset al . 1993; Li et al . 1993; Hacker et al . 1998; Jahn
1998). The presence of Triassic flysch and Jurassicmolasse along the suture also supports earlyMesozoic collision of the continental blocks (Li1996). However, K–Ar ages of phengite and horn-blende in the UHP metamorphic rocks are scat-tered widely , from the Proterozoic to the Mesozoic(Ishiwatari et al . 1990b; Li et al . 1994), possibly
due to excess argon inherited from their Precam-brian protoliths (Giorgis et al . 2000). The Chinesesuture is postulated to extend into the Korean Pen-
insula, namely into the Imjingang or Ogcheon belt(Ernst & Liou 1995; Ree et al . 1996), and further
continuing into the Hida (marginal) belt in Japan(Isozaki 1996, 1997). The latter idea is based onHiroi’s works of 1981 and 1983 (Hiroi 1981, Hiroi1983), which first correlated the Unazuki meta-morphic rocks in the Hida belt to the Ogcheon andImjingang (Y onchon or Y eoncheon) metamorphicbelts. In their model, the northern part of theKorean Peninsula belongs to the Sino-Korean
block, whereas its southern part belongs to theYangtze block. Kim et al . (2000) reported a Rb–Sr
mineral isochron age of 226 ±1.2 Ma for the mylo-
nite in the Gyeonggi (Kyonggi) massif on the southof the Imjingang belt, and interpreted it to repre-sent post-collisional, extensional ductile shear. Leeet al . (2000) correlated the early Proterozoic gran-
ulites of the Gyeonggi massif to that of the Y angtzecraton on the basis of zircon–monazite sensitivehigh mass-resolution ion microprobe (SHRIMP)U–Pb age and an Nd–Sm isotope model age (T
DM),
although they admit ‘it is probably not warrantedto attempt any tectonic correlation solely based onthe resemblance in T
DM ages or SHRIMP zircon
ages’.
The Imjingang metamorphic belt is a typical
Barrovian kyanite–sillimanite-type metamorphicbelt (Y amaguchi 1951). Ree et al . (1996) reported a
Nd–Sm mineral isochron age of 249 ±31 Ma, and
0.8–1.3 GPa and 630–790 ∞C metamorphic condi-
tions for the adjacent garnet amphibolite unit onthe south, and correlated it to the Permo-Triassicsuture in China, regarding the garnet amphiboliteas retrograded from UHP eclogite. However, ‘crit-ical evidence of UHP metamorphism such as eclog-ite, diamond and coesite remains to be found’ (Reeet al . 1996). Min and Cho (1998) identified a three-
stage metamorphic evolution of the Ogcheon belt:(1) Siluro-Devonian medium-pressure (0.5–0.8 GPaand 520–590 ∞C) metamorphism; (2) Triassic
regional retrograde metamorphism (0.1–0.3 GPaand 350–500 ∞C); and (3) Jurassic–Cretaceous ther-
mal metamorphism around granitoids. This meta-morphic history coincides with the structuraldevelopment of the Ogcheon belt as an early Pale-ozoic intracontinental rift zone evolved into anintracontinental fold–thrust belt without ophiolite(Cluzel et al . 1990), but is not consistent with the
Permo-Triassic intercontinental collision process
that involves HP metamorphism (Ernst & Liou1995). Moreover, another medium-pressure meta-morphic belt with a muscovite
40Ar–39Ar age of 200–
230 Ma is reported from the Fangshan area in the
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Paleozoic ophiolites and blueschists 199
Western Hills of Beijing (Fig. 6), that is, in the mid-
dle of the Sino-Korean craton (Wang & Chen 1996).As aforementioned, there is no direct evidence ofUHP–HP metamorphism in the Korean metamor-phic belts; it is likely that the 200–250-Ma medium-pressure metamorphic belts such as Imjingang,Ogcheon and Fangshan took place in the intracon-tinental fold–thrust belts of the Sino-Korean cratonduring the Y angtze–Sino-Korean collision.
The most convincing criteria with which to
define tectonic affiliation of an area may be stratig-raphy of the sediments covering continental base-ment. The Paleozoic system of the Ogcheon zoneshows typical ‘Sino-Korean’ stratigraphy charac-terized by thick Cambro-Ordovician limestone,late Paleozoic coal-bearing sediments, and ‘thegreat hiatus’ in between (Fig. 7). This is differentfrom the Y angtze stratigraphy with thick Siluro-Devonian shale (Fig. 7). This suggests that nomajor suture exists between the northern andsouthern parts of Korea. Lee et al . (1998) also state
that the Korean Peninsula as a whole belongs tothe Sino-Korean craton at least from the late Pro-terozoic in view of the overall similarity in age,geology , petrography and geochemistry .YAEYAMA PROMONTORY HYPOTHESIS
In view of the eastward-convex, winding geological
structure over the Korean Peninsula, Teraoka et al .
(1998) proposed that the Sulu suture does notextend to Korea, but turns southward beneath theYellow Sea. They did not specify , however, where
the destination of the redirected suture is. Theirintensive studies on the chemistry of clastic gar-nets in the Japanese Cretaceous–Tertiary sedi-ments indicate that these eclogitic, pyrope (Mg)-rich and spessartine (Mn)-poor garnets occur insandstones of the Shimanto accretionary complexin the outer zone of the South-west Japan andRyukyu arcs (Takeuchi 1992; Teraoka et al . 1999),
whereas such garnets are not found from the Cre-taceous–Paleogene fore-arc and intra-arc basins ofsouth-western Japan. This finding suggests thatthe Sulu UHP belt does not extend to Korea andnorth-eastern China, but turns to the south into theprovenance area of the Shimanto sediments. Wepropose that the Sulu suture reappears at IshigakiIsland, Y aeyama Archipelago, southern Ryukyu,and continues into Japan and Primorye, detouringaround Korea (Fig. 6). Along this highly sinuous
Fig. 6 Proposed sinuous configuration of the
eastern elongation of the Sulu–Dabie suture ( ca
250 Ma) of China passing subduction zones ofRyukyu, south-western Japan and Russian Pri-morye but detouring around Korea. A prelimi-nary version of this diagram appeared inIshiwatari and Tsujimori (2001).
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200 A. Ishiwatari and T. Tsujimori
suture line, continental collision took place in the
Chinese segment, whereas subduction of oceaniclithosphere took place in the Japanese–Russiansegment. This agrees with the argument of Ernstand Liou (1995) that UHP metamorphism takesplace in the continental collision segment, whilenormal HP metamorphism takes place in the oce-anic subduction segment, along a single suture line.
Such a sinuous configuration of the collisional
suture is also observed in the Alpine chain in theMediterranean area. It is noteworthy that the coes-ite-bearing UHP metamorphic rocks of the DoraMaira massif (Chopin 1984) occur at the north-western tip of the acute Adriatic (Apulian) prom-ontory , whose profile is clearly visible on thepresent-day seismic map (Mueller 1989). Ourmodel suggests that the Chinese UHP rocks occurat the southern tip of the Dabieshan promontory ofthe Sino-Korean Craton and at the northern tip ofthe Sulu promontory of the Y angtze craton. In thiscontext, the early Mesozoic HP schists of the Ish-igaki Island possibly represent the southern tip ofanother promontory of the Sino-Korean Craton,which we call Y aeyama promontory after theregional name for the southernmost RyukyuIslands. The late Paleozoic–early Mesozoic (200–
250-Ma) HP metamorphic belts in Ryukyu, Japanand Russian Primorye are suitable as an easternextension of the Chinese collisional suture of thesame age. The late Paleozoic–early Mesozoic(Indosinian) dextral ductile shearing reported fromthe Ogcheon belt (Cluzel et al . 1991; Otoh et al .
1999) is compatible with the reciprocal movementbetween the Sulu and Y aeyama promontories.
The Tananao schist complex of eastern Taiwan
has long been regarded as a late Paleozoic oro-genic belt (Fig. 4 of Cluzel 1991), but hornblendeK–Ar ages of the schist are younger than 90 Ma,and the associated gneiss and granite also give90 Ma or younger Rb–Sr and U–Pb ages (Yuliblueschist is as young as 10 Ma; Jahn et al . 1986).
This indicates that the late Paleozoic–early Meso-zoic metamorphic belt around the Y aeyama prom-ontory does not extend to Taiwan.
The Y aeyama promontory hypothesis provides
some insights for pre-Cretaceous paleogeographyof Japan. Paleozoic and Mesozoic fossil faunas indi-cate that the South Kitakami and KurosegawaBelts were situated further to the south of the Hidamarginal belt before the Late Cretaceous large-scale strike–slip movement (Otoh & Sasaki 1998;Tazawa 2000). These three belts show overall strati-
graphic similarity with the Khanka massif; theseterranes may have together developed along theactive continental margin of the Sino-Korean cra-ton, although the South Kitakami and KurosegawaBelts were later displaced toward the north bya Cretaceous left-lateral strike–slip movement(Tazawa 2000). Isozaki (1997) proposed thatJapanese accretionary complexes, including theSouth Kitakami block and ‘Y akuno oceanic plateau’,developed along the Y angtze continental margin,assuming that the Sino-Korean–Y angtze suturezone passes through central Korea and extends tothe Hida Mountains in south-western Japan. Asmentioned earlier, however, the fossil fauna andlithology of the Permo-Triassic cover of the Y akunoophiolite closely resemble those of Primorye(Nakazawa 1958). Permian strata of the Hida mar-ginal belt and South Kitakami belt also show a fau-nal kinship with those of north-eastern China andPrimorye (Tazawa 1993, 2000; Otoh & Y anai 1996;Otoh & Sasaki 1998). Our model infers that allJapanese Paleozoic terranes, except for accretedseamounts, have developed along the Sino-Koreanmargin. The South Kitakami and KurosegawaBelts may have been placed somewhere betweenKyushu and Ishigaki Island along the eastern mar-
gin of the Y aeyama promontory (Fig. 6).
Fig. 7 Schematic stratigraphic columns for the contrasting Paleozoic
sequences on the (a) Yangtze block (eastern Sichuan, lower Yangtze Val-ley) and (b) Sino-Korean block (Hebei, Shanxi; after Willis & Blackwelder1907). The paleozoic sequence in southern Korea is also of the Sino-Korean type.
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Paleozoic ophiolites and blueschists 201
PALEOZOIC OPHIOLITES IN JAPAN AND
PRIMORYE: GEOTECTONIC IMPLICATIONS
Cluzel (1991) proposed that the Y akuno ophiolite
formed in a rift zone in the Sino-Korean continen-tal margin in Middle–Late Carboniferous time,and was emplaced onto the ‘Honshu block’ in themiddle Permian by the closure of the small seabasin between the rifted continental blocks. Such a‘Tethyan ophiolite’ model including continental
rifting, sea-floor spreading, and obduction may beapplicable to the Cambrian Khanka ophiolite inPrimorye, which may have formed before thebeginning of the circum-Pacific-type orogeny , butmay not apply for the other circum-Pacific ophio-lites (Ishiwatari 1994). The Y akuno ophiolite is tec-tonically underlain by the Permian Ultra-Tambaaccretionary complex, which is in turn underlainby the Jurassic Tamba accretionary complex. Eachof the two complexes consists of ‘oceanic platestratigraphy’ (Isozaki 1996); successive underplat-ing of oceanic and trench-fill sediments beneaththe Y akuno ophiolite may have developed theseaccretionary complexes. The Miyamori ophiolitealso thrust over the Jurassic accretionary complex(Tazawa 1988). These ophiolites did not thrust ontoold continental blocks as imagined in the Tehyanmodel. Instead, younger accretionary complexesformed beneath the old ophiolites.
In contrast, Isozaki (1996, 1997) proposed that
the Y akuno ophiolite with thick crust representsoceanic plateau, which has formed in the midst ofthe ocean by a superplume activity and lateraccreted to Japan. However, the sedimentarycover of the Y akuno ophiolite is thick black shalewith a restricted age of radiolarian fossils (middlePermian), which is incompatible with a long plate-
tectonic travel in the ocean.
Ishiwatari et al . (1990a) divided the gabbroic
rocks of the Y akuno ophiolite into two types: aMORB type in the eastern area and an island-arcbasalt (IAB) type in the western area, according tothe chemistry of coexisting clinopyroxene and pla-gioclase. They postulated the Y akuno ophiolite asbeing a cross-cut section of oceanic island arc andan adjacent back-arc basin, which was affected bya mantle plume. The granulite-facies metacumu-late represents the basal part of thickly developedmafic crust of the island arc and back-arc basin.The mantle section of the Miyamori ophiolite isalso interpreted as hydrous mantle beneath islandarc (Ozawa 1988). Mantle peridotite of the Oeyamaophiolite resembles that from the ocean floor, andbears some podiform chromitite with hydrous min-eral inclusions such as Na-phlogopite and par-
gasite (Matsumoto et al . 1995), suggesting either a
supra-subduction zone (SSZ) or a fast spreadingridge setting (Arai 1997). However, the SSZ set-ting is more preferable in view of the orbicularchromitite with a very high Cr# (0.76–0.85)reported by Y amane et al . (1988).
It should be noted that the ophiolite sequence is
actually exposed in the submarine trench wallsaside the Mariana arc (Bloomer & Hawkins 1983)and Tonga arc (Bloomer & Fisher 1987; Fig. 8). Itis important that back-arc spreading is activebehind these island arcs, and the accretionarycomplex is currently absent in the subductionzone. Moreover, lawsonite-bearing blueschistblocks were drilled from serpentinite seamounts(diapirs) in the Mariana fore-arc area (Maekawaet al . 1993, 1995; Fig. 8). The ophiolite–blueschist
association is well demonstrated in the Japan–Primorye area such as the Oeyama ophiolite–Renge blueschist, Y akuno ophiolite–Suoblueschist, Miyamori ophiolite–Motai blueschistand Sergeevka ophiolite-Shaiginskiy blueschist. Incontrast, subduction zones such as the JapanTrench and Nankai Trough have formed huge
accretionary complexes from the Cretaceous tothe present (Taira 1985; Fig. 8).
It is likely that the accretion period and non-
accretion period as represented by the presentNankai Trough and Mariana Trench, respectively ,have repeated one after another and from segmentto segment in the history of the Japan–Primoryeaccretionary orogenic belt. The ophiolite–blue-schist period with tectonic erosion at the subduc-tion zone may have been followed by a period ofmassive accretion. This idea is quite compatiblewith the geochemical island-arc and marginal-basin signatures of the ophiolitic rocks.
CONCLUSIONS
Age, lithology , and structural position of the Pale-
ozoic ophiolites and blueschist in south-westernJapan and Primorye strongly support original con-tinuation of the geological units over the two sidesbefore the Miocene opening of the Japan Sea. Thenewly obtained K–Ar ages of the blueschists andassociated gneiss complex of Primorye (250 Ma)coincide with the widespread metamorphic eventsin East Asia such as the UHP metamorphism of theSulu–Dabie suture and medium- and low-pressuremetamorphism in Japan (Unazuki and Hida), Korea(Imjingang and Ogcheon), and northern China
14401738, 2003, 2, Downloaded from https://onlinelibrary.wiley.com/doi/10.1046/j.1440-1738.2003.00390.x by Ohio State University University Libraries, Wiley Online Library on [10/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
202 A. Ishiwatari and T. Tsujimori
(Fangshan). The wide-scale convergence resulted
in the UHP metamorphism along the continentalcollision zone, but ordinary HP metamorphism tookplace along the sinuous oceanic extension of thesame suture passing Ryukyu, Japan and Primoryebeyond the Y aeyama promontory , detouring aroundKorea. Although 250-Ma HP schists are rare in
Japan, slightly older (Renge) and younger (Suo)blueschists are preserved in many places, suggest-ing persistent subduction. The repeated formationof the ophiolite–blueschist assemblages and the tec-tonically underlying, younger accretionary com-plexes suggests repetition of the non-accretingsubduction as in the present Mariana Trench andthe accreting subduction as in the Nankai Troughthrough development of the orogenic belts in Japanand the Russian Far East.
ACKNOWLEDGEMENTS
We thank Professors A. I. Khanchuk, S. V .
Vysotskiy , S. A. Shcheka and S. V . Kovalenko of the
Far East Geological Institute of the Russian Acad-
emy of Sciences in Vladivostok for their coopera-tion in the field and sample preparation. ProfessorTetsumaru Itaya of Okayama University of Sci-
ence is thanked for providing equipment andinstructions for K–Ar dating for the second author.
Professors J. G. Liou and H. Maekawa arethanked for constructive reviewing of the manu-script. The first author acknowledges Grant-in-Aidfor Scientific Research (C)-(2)-10640462 and-14540447 provided by the Ministry of Education,Japan. The second author acknowledges Grant-in-Aid for JSPS Fellows.
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Lithos 342-343 (2019) 345-360 Contents lists available at ScienceDirect Lithos ELSEVIER Zircon U-Pb age and Nd isotope geochemistry of latest Neoproterozoic to early Paleozoic Oeyama ophiolite: Evidence for oldest MORB-type oceanic crust in Japanese accretionary system and its tectonic implications Kosuke Kimura *, Yasutaka Hayasaka Department of Earth and Planetary Systems Science,Hiroshima University,Higashi-Hiroshima 739-8526,Japan ARTICLEINFO ABSTRACT Article history: The Oeyama ophiolite is considered as one of the major ophiolites in the Japanese Islands that forms a part of the Received 15 January 2019 "circum-Pacific Phanerozoic multiple ophiolite Belts" In this study, zircon U-Pb age and their REE geochemistry, Received in revised form 25 May 2019 Accepted 3 June 2019 precise crystallization age and characterize the proper tectonic setting of this ancient oceanic crust. The zircon Available online 6 June 2019 grains of the two studied gabbroic samples yield 206pb/238U ages of 545.4 ± 2.6 Ma and 532.4 ± 3.1 Ma. From five gabbroic rocks, a comparable 566 ± 95 Ma Sm-Nd whole rock isochron age is obtained. The time- Keywords: Oeyama ophiolite corrected eNd values (t = 545 Ma) for each sample are ranging from +7.2 to +7.7 that suggest their normal Southwest Japan mid-oceanic basalt (N-MORB) affinity. Discrimination diagrams using zircon U/Yb versus Hf and Y also indicate Zircon U-Pb age their oceanic crustal origin. This is so far known the oldest date of ophiolite in the Japanese accretionary system. Sm-Nd isotopes On the basis of the present dataset, we have offered a scenario of the potential tectonic development highlighting Zircon chemistry the journey of this group of oceanic crustal rocks from the mid-oceanic region till their final accretion to the Japanese Islands. The present tectonic model envisages an initial oceanward subduction at ca. 530 Ma switched over to the present continentward subduction at ca. 400 Ma. Moreover, the three different units of the Oeyama ophiolite have been re-classified based on the present age dataset, and the formation and evolution of these units constituting the typical lower part of the ocean plate are explained. 2019 Elsevier B.V. All rights reserved. 1. Introduction oceanic crust and mantle evolution of the paleo-Pacific region during the Phanerozoic time. Numerous studies have been carried out on ophiolites to unravel the One of the important characteristics of the ophiolite occurrences formation and evolution of the ancient oceanic crust and chemical evo- forming the “circum-Pacific Phanerozoic multiple ophiolite belts" is lution of the mantle (e.g. Hopson et al., 1981; Ichiyama et al., 2008; the ocean-ward younging age polarity, implying the continuous accre- Ishiwatari, 1985; Ishiwatari and Tsujimori, 2003; Mankinen et al., tion of oceanic materials through Phanerozoic (Ishiwatari, 1994). 2002; Sano et al., 2000; Tatsumi et al., 2000). In the circum-Pacific re- Away from those occurring in the Japanese Islands, several ophiolite gion, accretionary complexes had been formed at different times in bodies in the Klamath Mountain (Latest Neoproterozoic to Early Paleo- the Phanerozoic, and as a result the fragment of oceanic crust of differ- zoic Trinity ophiolite to Late Jurassic Josephine ophiolite) are juxtaposed ent ages and origins are involved in the crustal growth in this region together with accretionary complexes and are younging towards the (e.g. Ichiyama et al., 2008; Ishiwatari, 1985; Ishiwatari and Tsujimori, Pacific Ocean (Harper et al., 1994; Ishiwatari, 1994; Wallin et al., 2003; Sano et al., 2000; Shervais et al., 2004; Tatsumi et al., 2000). 1995; Wallin and Metcalf, 1998). Furthermore, similar accreted ophio- These fragments of oceanic crusts represented by the ophiolite bodies lite sequence is also suggested from the Far-East Russia and around are coined as “circum-Pacific Phanerozoic multiple ophiolite belts" by the Eastern Australian region (Ishiwatari et al., 2003; Lus et al., 2004). Ishiwatari (1994) on the basis of their age and distribution. Therefore, Hence, detailed geochronological data of formation of such ophiolites it is widely believed that such ophiolites are useful to study the ancient in Japanese Islands is considered to be important to regional age-wise correlation as well as their broad status in the global occurrence of "circum-Pacific Phanerozoic multiple ophiolite belts". * Corresponding author. Ophiolitic rocks of Japanese Islands are well-compiled by Ishiwatari E-mail address: kimurakoske@hiroshima-u.ac.jp (K. Kimura). et al. (2016). Major ophiolitic rocks in Japan can be broadly grouped https://doi.org/10.1016/jlithos.2019.06.001 0024-4937/@ 2019 Elsevier B.V. All rights reserved. 346 nura,Y.Hayasaka/Lithos342-343(2019)345-360 a) (b) InnerZone Nm TTL ISTL S-R M-S Yk MTL UYTL MTL Kr ophioliticrocks ISTL Mk Tectonic Line — Early Paleozoic OuterZone .-- Mesozoic UYTL LatePaleozoic ..Cenozoic Fig. 1. (a) Distribution ofthe ophiolitic rocks in the Japanese Islands. S-R: Sangun-Renge Belt, M-H: Miyamori-Hayachine ophiolite, Mt: Motai metamorphic rocks, Kr: Kurosegawa Tectonic Belt, Yk: Yakuno ophiolite, Mk: Mikabu Belt, Hr: Horokanai ophiolite, Km: Kamuikotan ophiolitic melange, Pr: Poroshiri ophiolite, M-S: Mineoka-Setogawa ophiolite. MTL: Median Tectonic Line, UYTL: Usuki-Yatsushiro Tectonic Line, TTL: Tanakura Tectonic Line, ISTL: Itoigawa-Shizuoka Tectonic Line. (b) Distribution of ophiolite bodies of Sangun-Renge elt, Southwest Japan. Ng: Nagato Tectonic Belt, Ym: Yamaga metagabbroic body, Nm: Nomo metagabbroic Complex. The marked box represents the distribution of the presently studied ophiolite bodies, the details are as in Fig. 2. according to their ages, e.g. (1) Early Paleozoic Oeyama ophiolite in the (4) Mesozoic Mikabu greenstones in the Mikabu Belt, and Horokanai Sangun-Renge Belt, Miyamori-Hayachine ophiolite and Motai meta- ophiolite, Kamuikotan ophiolitic melange and Poroshiri ophiolite dis- morphic rocks in the South Kitakami Belt, and some ophiolitic rocks in tributed in Hokkaido, and (5) Cenozoic Mineoka-Setogawa ophiolite the Kurosegawa Tectonic Belt, (2) late Paleozoic Yakuno ophiolite in (Fig. 1). The above-mentioned ophiolite units in the Japanese accretion- the Maizuru Belt, (3) late Paleozoic to early Mesozoic greenstones ary system also have a Pacific sea-ward younging age polarity, as a (ophiolitic origin) in the Paleozoic to Mesozoic accretionary complex, whole. For example, in the Inner Zone of Southwest Japan, piled nappe (a) lzushi Wakasa Oeyama Sekinomiya Ashidachi (Fig.2b) Tari-Misaka Ultramafic rocks Osayama Mafic suite Saijo 30km (Fig. 2c) (b) LEGEND C Japan. (b) Detailed geological map of the Ashidachi body of the Oeyama ophiolite (modified after Hayasaka et al, 1995). (c) Detailed geological map of the Saijo body of the Oeyema ophiolite. Sampling points are noted in the map. K.Kimura, Y. Hayasaka / Lithos 342-343 (2019) 345-360 347 Table 1 Sample LG1 PI GB3 PI GB3 Cpx Fine-grained gabbro Cpx No 1 2 3 1 2 3 1 2 3 1 2 SiO2 53.35 53.16 53.39 51.18 52.02 52.57 51.56 51.01 49.45 50.76 50.23 50.80 TiO2 11.93 11.70 11.66 13.32 12.90 12.40 21.44 21.71 20.10 0.86 0.99 0.53 Al203 0.22 0.25 0.25 0.39 0.47 6.81 serpentinized ultramafic rocks form two wide zones; Renge serpentinite (sp.) zone and Happo sp. zone (modified Suda et al., 2014a). (b) 6.89 4.55 5.15 4.47 29.07 Fe0 28.86 28.84 29.79 29.76 29.52 2.56 3.55 4.37 7.39 6.51 5.15 MnO 0.02 0.00 0.06 0.06 0.02 0.09 0.59 0.86 0.52 0.20 0.16 0.15 MgO 0.03 0.06 0.00 0.02 0.02 0.00 0.21 0.14 0.19 16.53 16.40 16.47 Cao 0.03 0.04 0.02 0.03 0.01 0.02 14.72 15.09 15.34 18.15 19.74 20.60 Na20 0.09 0.10 0.08 0.06 0.07 0.06 0.00 0.00 0.01 0.47 0.26 0.31 then immediately sealed with aluminum crimp seals. In a 0.00 0.00 0.00 0.01 0.01 000 60'0 0.25 0.24 0.11 0.00 0.00 Cr203 4.70 4.79 5.00 3.90 4.13 4.55 0.36 0.49 0.30 0.06 0.25 0.98 total 99.44 98.94 99.29 98.76 99.40 99.53 98.33 99.75 97.41 99.08 69'66 99.46 0=24 !S 7.28 7.29 7.30 7.08 7.13 7.19 7.75 7.57 7.50 7.52 7.41 7.49 Ti 1.75 1.72 1.71 1.97 1.90 1.82 3.45 3.45 3.27 0.10 0.11 0.06 Al 0'0 80'0 80'0 0.05 0.05 0.04 0.86 0.83 0.87 0.79 0.89 0.78 Fe 4.68 4.67 4.65 4.85 4.81 4.76 0.45 0.62 0.78 0.92 0.80 0.63 Mn 0.00 0.00 0.01 0.01 0.00 0.01 0.07 0.10 0.06 0.03 0.02 0.02 Mg 0.00 0.01 0.00 0.00 0.00 Fig. 1. Geological map of western Shiroumadake area. The star symbols represent the location of the Hakuba Happo hot spring. (a) The K. Suda et al. / Geochimica et Cosmochimica Acta 206 (2017) 201-215 0.02 0.02 3.65 3.60 3.62 Ca 0.01 0.01 0.00 0.01 0.00 0.00 3.30 3.34 3.47 2.88 3.12 3.25 PN 0.01 0.02 0.01 0.01 0.01 0.01 0.00 0.00 0.00 0.14 0.07 0.09 K 00'0 0.00 0.00 0.00 0.00 0.00 0.01 0.03 0.03 0.02 0.00 0.00 Cr 1.25 1.28 1.32 1.05 1.10 1.21 0.11 0.14 0.09 0.01 0.03 0.11 15.00 15.02 15.04 15.02 15.01 15.03 16.01 16.08 16.09 16.06 16.06 16.05 An 0.58 0.57 0.56 0.65 0.63 0.60 Wo 45.6 46.9 44.4 39.6 42.4 44.3 En 43.5 45.3 47.2 50.2 49.1 49.2 Fs 10.9 7.8 8.4 10.3 8.5 6.5 bodies including ophiolitic rocks are showing younging direction (from oldest units of the accreted component in Japanese Islands (Ishiwatari uppermost unit downwards) while considering together with the asso- and Hayasaka, 1992; Ishiwatari and Tsujimori, 2003). Hence, this partic- ciated accretionary complexes and high P/T metamorphic rocks ular ophiolite body becomes immensely important to understand the (Hayasaka, 1987). Among them, the Oeyama ophiolite occupies the up- early history of the Japanese Islands and the geochemical characteriza- permost nappe of the Inner Zone of Southwest Japan, and is one of the tion of the Paleo-Pacific oceanic plate. a b) Fig.3.Outcrop-scale features of the (a) coarse-grained diallage gabbro (GB3) with 2 cm coin; (b) coarse- grained leucogabbro (LG1) and (c) the contact relationship between GB3 and LG1. Note the formation of chilled margin. 348 K. Kimuro, Y. Hayasaka / Lithos 342343 (2019) 345360 Reliable high precision age data and the understanding of the tec- is the first successful separation of zircon from the rocks of the Oeyama tonic settings of the Paleozoic ophiolites in Japan are scarce. This is par- ophiolite. We measured their U-Pb ages and chemistry using laser- ticularlysofortheeyamaphioliteasthereliablehighprecisionae ablation inductively coupled mass spectrometer (LA-ICP MS). Further- data of it are still lacking. This leads to the limited understanding of more, Sm-Nd isotopic data were generated and also processed based u anedodde, a pue poq s o ssaod uoeog a on this newly obtained zircon age. perspective of the regional tectonic development of accretionary com- plex in the proto-Japan. This holds the key also to have a finer under- 2.Geologicalbackground standing of the global-scale correlation and tectonic history of the neighbors of Paleozoic supercontinent, “Gondwana". The Oeyama ophiolite is one of the constituent units of the Sangun- Hence, in this study zircon U-Pb geochronology, Nd-Sr isotopic com- Renge Belt in the Inner Zone of Southwest Japan (Fig. 1) covering all positions and zircon chemistry for the mafic rocks were aimed. The high other units as a nappe in the belt. According to Kurokawa (1985), the spatial resolution zircon U-Pb age dating can record high precision crys- peridotite body located at Mt. Oeyama, northern Fukuchiyama City, tallization age which are difficult to get reset in the subsequent low Kyoto Prefecture is one of the type sections (Fig. 2a). The studied temperature metamorphism and deformation. Similarly, zircon chemis- ophiolite bodies are presently distributed towards the WSW direction try and Nd-Sr isotopic composition are also considered to be robust of the Oeyama peridotite body. in the Izushi, Sekinomiya, Wakasa, eep asau le pnss juasaid au u saueo juanbasqns Aue ysuege Osayama, Ochiai-Hokubo, Ashidachi (Fig. 2b) and Tari-Misaka areas, are used to clarify the high precision crystallization age and to evaluate and finally can be traced to the Saijo body in northeastern Hiroshima the tectonic setting of the Oeyama ophiolite. Zircon grains are found Prefecture (Fig. 2c). Moreover, in the further western areas mafic- from two samples of the studied coarse-grained diallage gabbro. This ultramafic rocks are exposed as a part of the Sangun-Renge Belt, namely a b CDX 2mm 2mm 2mm 2mm 0.5mm 0.5mm Fig 4. (a) Photomicrograph of coarse-grained diallage gabbro (GB3), open polar, (b) the same under cross polar: (c) Photomicrograph of coarse-grained leucogabbro (LG1), open polar, (d) the same under cross polar. K.Kimura,Y.Hayasaka/Lithos 342-343(2019)345-360 349 Nagato Tectonic Belt of Yamaguchi Prefecture, Wakamiya and Sasaguri tinite-hosted hydrothermal systems; Lost City field orite, one of the carbonaceous chondrite group meteorites, Tari-Misaka body of the Oeyama Ophiolite. area of Kumamoto Prefecture and Nomo metagabbroic Complex at The meta-cumulates of the Unit II occur in the Oeyama and Wakasa Nagasaki Peninsula (Isozaki and Tamura, 1989; Nishimura and body, and also Wakamiya-Sasaguri area in the western extension of the compositional features of C, to Cs n-alkanes for Hakuba Sangun-Renge Belt, and had suffered the epidote-amphibolite-facies from sedimentary and felsic volcanic rocks are believed to metamorphism. The meta-cumulate of the Unit Il is suggested to be Sherwood Lollar et al., 2002). Abiotic synthesis has been the tectonic block that derived from troctolite to anorthosite and differs that containing granitoids are considered to be a separate lithological bution cannot be identified as definitely being of abiogenic ethane (C2), propane (Cs), iso-butane (i-C4), n-butane (n-C4), iso- The metamorphic grade of the Unit Ill is lower than that of the unit II. ern extension of the Oeyama ophiolite body can also be traced to- show a similar isotopic trend (Yuen et al., 1984). These 5.1.1. Comparison with potentially abiogenic hydrocarbons in has been observed in n-alkanes for the seafloor serpen- area of Fukui Prefecture, Naradani and Fukuji areas of Gifu Prefecture, fered low grade metamorphism (Unit III). and Happo-one, Kotaki, Renge and Omi areas spreading in Nagano, The present study focusses on the Ashidachi body and the Saijo body at the geologic setting that is not directly associated with of the Oeyama ophiolite. The former body occurs in the northwestern tion related metamorphism is evidenced by the occurrence of jadeitite Okayama prefecture. This one overlies the Permian accretionary com- and blueschist rocks in the melange associated with the Sangun- plex of the Akiyoshi Belt as a low-angle nappe, and has mafic rocks of Renge Belt (Tsujimori and Itaya, 1999) occurring below the nappe of the Unit Ill that spread over a relatively wider area than the other the Oeyama ophiolite. Oeyama ophiolite bodies. The Unit Il is conspicuously absent in the carbons (C2—Cs) with respect to Ci in natural gases Ashidachi body. Ultramafic rocks of the Unit I are mainly composed of three units (Kurokawa, 1985): Unit I (82 vol%) is ranging from ultrama- dunite, harzburgite and lherzolite, and are geologically comparable to fic tectonite to lowest cumulate unit that consists of dunite, (lherzolitic) the residual peridotite of upper mantle (Hayasaka et al., 1995). These (Proskurowski et al., 2008) and Logatchev-II field typical trend for hydrocarbon gases produced by the mafic metamorphosed cumulate. Unit Ill (3 vol%) is weakly metamor- to Cs alkanes from Hakuba Happo hot spring. Happo #1 samples phosed (up to greenschist facies) coarse-grained diallage gabbro and di- (Charlou et al., 2010) (Fig. 4a). On the other hand, even Proskurowski et al., 2008; Wang et al., 2009; Charlou ber 2013 (square), January 2014 (diamond), October 2014 (circle), Comparison of concentrations of dissolved gases and organic acids. (e.g., Guo et al., 1997). Moreover, the Ci—C4 n-alkanes between the 813C value and the inverse carbon number is (Yuen et al.,1984; Sherwood Lollar et al., 2002; methane concentration is elevated compared to the pre- younger gabbro is the fine-grained hornblende-gabbro that intruded spinel: Ishiwatari, 1991b) is contained characteristically in harzburgite inverse carbon numbers of Ci—Cs n-alkanes, is found in Table 2 Whole-rock chemistry of coarse-grained diallage gabbro and leucogabbro from Ashidachi and Saijo bodies Ashidachi body Saijo body 5-19-4 5-20-3 5-20-9 5-25-10C GB2 GB4 92061305 GB3 LG1 SiO2 48.86 47.21 Lost City_IF/2 48.61 49.25 46.39 50.83 51.04 52.25 TiO2 1285 0.23 0.22 SE0 0.34 1373 0.54 0.38 0.25 VPDB 14.94 19.74 16.43 14.58 17.14 18.88 18.09 Table 2 813C Fe203 4.44 3.54 4.37 4.16 104 5.65 -35 6.42 3.09 MnO 10.7 0.08 1250 3720 0.12 0.11 0.12 0.14 0.09 MgO 1815 6.62 10.21 10.38 -45 10.36 field 9.81 3.16 Cao 16.26 16.03 25 15.83 6±0.2 13.46 11.32 12.67 10.23 Na20 1.69 2.31 1.81 1.86 2.86 -40 3.85 2.86 5.15 K20 0.13 0.27 0.23 0.44 0.26 0.77 0.44 0.16 0.29 P205 Happo #1 n-alkanes show an isotopic depletion in 13C 0.009 A similar positive linear trend on the 8l3C vs. (1/n) plots 0.012 0.012 0.010 0.013 Acetate (μmol/L) Lost Total 96.28 96.04 96.50 96.32 98.18 97.46 97.29 pentinite-hosted 98.16 Sc 一一一一 38.4 664 48.3 17.2 V 158.0 121.1 172.5 66.4 Cr 678.2 1000.7 381.2 72.6 30.2 44.2 36.1 15.4 Ni 147.8 223.3 194.7 195.5 130.7 244.8 96.9 121.4 31.7 Cu 133.7 80.5 not just a local characteristic for Happo #1 site, implying 10.8 1936; 103.5 10.7 decomposition 55.6 Zn 31.4 26.8 30.0 32.2 34.7 35.8 47.5 64.6 46.3 Ga 一 - 12.9 12.8 9.2 16.9 Rb 3.8 6.2 6.0 10.8 6.5 16.4 35.6 6.8 11.1 Sr 161.5 Formate (μmol/L) C2H6 (μmol/L) -30 tem occurs throughout the world. gases from commercial wells located in the Xujiaweizi and 813C3). Such a 13C depletion trend is the opposite of a 174.1 225.3 Y 11.4 8.4 9.8 12.0 9.7 6.7 Vs. 12.2 8.3 Zr 14.6 15.1 9.1 21.3 6.5 2.9 15.7 11.0 12.5 Nb 0.6 1.8 0.6 <LLD 0.9 1.2 <LLD 1.0 1.5 Cs 一 0.7 <LLD <LLD 0.5 Ba 61.9 115.9 25.6 34.7 La 124 <LLD 0.0 1.1 Ce 0.8 0.4 2.0 2.8 PN 3.1 0.3 0.6 2.7 Yb 1.4 1.1 0.5 1.0 JH <LLD <LLD <LLD <LLD W <LLD <LLD <LLD <LLD Pb 3.5 0.5 1.8 2.4 0.1 1.2 5.9 4.0 2.1 Th 3.4 1.5 2.3 1.8 5.4 5.0 2.6 2.3 1.4 U <LLD <LLD 1.1 <LLD 一 一 LLD: lower limit of detection. 350 K.Kimura,Y.Hayasaka/Lithos 342-343(2019)345-360 -40 field —LG1 of City -Ashidachi jdu 0. ies 0.01 RbBaNbKLaCeSrPNdZrTiY VScMnCrNi -45 -GB3 1250 3720 Location Lost Cityb 10 nple m D es 0.01 Fig. 5. Spider diagrams. (a) primitive-mantle normalized plot of coarse-grained gabbro samples. Values of primitive mantle was from Lyubetskaya and Korenaga (2007), (b) global MORB normalized plot of the same gabbro rocks. The values of global MORB were from Arevalo and McDonough (2010). (Hayasaka et al., 1995). The older gabbro and the younger gabbro can be Prefecture. The ultramafic rocks also suffered serpentinization, and con- correlated with diallage gabbro and diabase of Unit Ill of the Oeyama tain anhedral vermicular Cr-spinel. This body also has two types of body, respectively. Metabasalt is rarely associated with the younger mafic rocks correlated with Ashidachi body, that is, older coarse- gabbro. These two types of gabbroic rocks also show different geochem- ical characters, the older gabbro has MORB affinity, whilst the younger and metabasalt. A small amount of leuco-gabbro occurs in association gabbro has the back-arc and/or arc-related features (Hayasaka et al., with the coarse-grained diallage gabbro. 1995). In the previous studies, almost all radiometric ages of the Oeyama The Saijo body that consists of the mafic-ultramafic complex similar ophiolite had been determined by K-Ar method for hornblende. to the Ashidachi body occurs in the Saijo area, northeastern Hiroshima Meta-cumulate of Unit Il yielded age of ca. 403-443 Ma from the (a) (b) 1370 Happo #1 4150 Table 2 1373 1285 813C 11140 6 ± 0.2 12260 Lost 9137 1815 26 ± 0.5 15kV X350 50um 15kV X250 10.7 Fig. 6. Cathodoluminescence (CL) images of zircon grains separated from (a) GB3 and (b) LG1. They show euhedral shape and igneous oscillatory zoning. K.Kimura,Y.Hayasaka/Lithos342-343(2019)345-360 351 Table 3 LA-ICP-MS U-Pb age data of zircon grains from the Saijo mafic-ultramafic body. Spot name 238U/206pb* ± 20 207pb*/206pb*± 20 206pb*/238U age± 20 (Ma) 207pb*/235U age ± 20 (Ma) 207pb*/206pb*age±20(Ma) Th/U Disc. % GB3 Coarse-Grained diallage gabbro 006SJGB3 11.27 ± 0.39 0.0648 ± 0.0045 548.2 ± 18.4 592.9 ± 35.2 767.7 ± 152.5 0.61 8.2 007SJGB3 11.23 ± 0.40 0.0575 ± 0.0067 550.0 ± 19.0 542.1 ± 52.1 509.1 ± 277.7 0.57 -1.4 008SJGB3 11.48 ± 0.38 0.0618 ± 0.0039 538.3 ± 17.1 563.5 ± 31.2 666.8 ± 141.0 0.95 4.7 009SJGB3 11.34 ± 0.35 ear correlation against the inverse carbon number when 544.8 ± 16.2 544.2 ± 35.6 the first C—C bond is formed during ethane formation. A 0.72 -0.1 010SJGB3 11.13 ± 0.36 0.0585 ± 0.0058 554.5 ± 17.0 553.0 ± 45.8 (Hu et al., 1998; Taran et al., 2007; McCollom et al., 0.72 -0.3 011SJGB3 11.32 ± 0.36 0.0584 ± 0.0056 545.5 ± 16.8 The other model has been presented by McCollom et al. 544.4 ± 224.9 0.64 0.0 016SJGB3 11.20 ± 0.35 0.0592 ± 0.0045 551.3 ± 16.4 555.6 ± 35.7 573.3 ± 174.5 0.78 0.8 106SJGB3 11.54 ± 0.37 0.0569 ± 0.0041 536.0 ± 16.5 526.9 ± 33.0 487.8 ± 167.3 0.55 -1.7 107SJGB3 11.54 ± 0.37 any isotopic fractionation. In contrast to the former model, 535.6 ± 16.5 554.2 ± 34.8 631.8 ± 165.5 0.51 3.5 108SJGB3 dized single-carbon substrates (e.g., CO, CO2, and 0.0605 ± 0.0047 528.1 ± 15.7 546.2 ± 35.7 622.3 ± 175.4 0.81 3.4 109SJGB3 11.71 ± 0.46 0.0578 ± 0.0063 528.4 ± 19.8 527.3 ± 48.9 precursors based on the assumption of negligible isotope 0.79 -0.2 110SJGB3 11.50 ± 0.40 0.0629 ± 0.0044 537.3 ± 18.1 570.6 ± 34.6 705.5 ± 156.4 0.55 6.2 111SJGB3 5.2.1. Derivation of an ideal equation based on compound K. Suda et al. /Geochimica et Cosmochimica Acta 206 (2017) 201-215 546.3 ± 15.7 546.7 ± 26.6 548.6 ± 122.5 0.54 0.1 116SJGB3 11.30 ± 0.33 0.0577 ± 0.0047 546.6 ± 15.2 541.3 ± 36.7 518.9 ± 188.4 0.74 1.0 117SJGB3 chains, resulting in more 12C incorporated into higher 0.0569 ± 0.0044 541.2 ± 16.6 531.3 ± 34.9 489.0 ± 179.5 0.62 -1.8 118SJGB3 11.27 ± 0.34 0.0599 ± 0.0040 fractionation due to a rapid reaction rate of polymeriza- 558.1 ± 31.8 599.4 ± 152.1 0.56 1.8 119SJGB3 11.20 ± 0.31 0.0601 ± 0.0041 551.3 ± 14.8 562.3 ± 32.5 607.2 ± 156.7 0.81 2.0 120SJGB3 11.49 ± 0.34 0.0610 ± 0.0040 538.0 ± 15.5 557.8 ± 31.4 639.4 ± 148.7 0.74 3.7 121SJGB3 11.48 ± 0.37 0.0589 ± 0.0047 538.6 ± 16.6 synthesis via polymerization 564.3 ± 184.5 0.80 0.9 126SJGB3 11.42 ± 0.34 0.0559 ± 0.0044 541.2 ± 15.6 523.8 ± 35.4 448.8 ± 185.8 0.79 -3.2 127SJGB3 11.54 ± 0.32 0.0586 ± 0.0043 535.6 ± 14.4 538.7 ± 33.6 551.8 ± 170.2 0.77 0.6 128SJGB3 12.37 ± 0.40 0.0632 ± 0.0045 501.0 ± 15.4 541.3 ± 33.3 714.8 ± 158.5 0.57 8.0 129SJGB3 11.68 ± 0.42 0.0578 ± 0.0073 529.5 ± 18.3 528.3 ± 55.4 523.5 ± 303.3 0.91 -0.2 130SJGB3 11.58 ± 0.31 0.0645 ± 0.0038 534.1 ± 13.9 578.7 ± 29.1 758.2 ± 129.7 0.64 8.4 131SJGB3 depleted in i3C relative to the reactant carbon source (Hu compounds, in common with the polymerization models 575.4 ± 14.9 709.7 ± 30.6 1161.5 ± 108.8 1.03 23.3 136SJGB3 11.33 ± 0.29 0.0576 ± 0.0028 545.2 ± 13.6 539.0 ± 23.3 513.0 ± 109.1 0.89 1.1 137SJGB3 11.08 ± 0.31 methane, while the C2--Cs alkanes showed a trend of con- 556.9 ± 15.0 562.7 ± 29.8 586.4 ± 140.6 0.65 1.0 138SJGB3 11.28 ± 0.33 0.0606 ± 0.0042 547.7 ± 15.2 563.0 ± 33.4 625.4 ± 158.6 0.75 2.8 139SJGB3 11.00 ± 0.29 0.0612 ± 0.0033 560.9 ± 14.0 additionof single-carbon 646.8 ± 120.5 0.83 3.1 140SJGB3 pound and C2+ hydrocarbons appear to show a positive lin- of single-carbon compounds (methylene units are assumed products have been reported by the experimental studies for 501.6 ± 30.9 499.0 ± 166.9 0.70 -0.1 141SJGB3 10.85 ± 0.28 0.0589 ± 0.0019 568.2 ± 14.2 567.2 ± 18.4 563.3 ± 73.6 0.71 -0.2 146SJGB3 11.27 ± 0.32 0.0570 ± 0.0033 548.1 ± 14.7 537.6 ± 27.1 493.3 ± 133.2 0.53 -1.9 147SJGB3 11.71 ± 0.40 account for observed carbon isotopic compositions in nat- ple. The McCollom et al. model assumes that the reversal 538.4 ± 47.5 582.1 ± 248.6 0.54 2.0 148SJGB3 11.03 ± 0.31 0.0596 ± 0.0036 559.4 ± 15.0 565.0 ± 29.0 587.8 ± 135.8 0.89 1.0 149SJGB3 11.42 ± 0.26 0.0583 ± 0.0016 541.2 ± 12.0 541.3 ± 15.2 542.0 ± 62.4 0.57 0.0 150SJGB3 11.23 ± 0.33 0.0579 ± 0.0043 549.7 ± 15.3 545.1 ± 33.9 526.0 ± 171.0 0.73 -0.8 151SJGB3 11.20 ± 0.28 0.0570 ± 0.0034 ues of C2—-Cs alkanes increase with increasing carbon num- 539.7 ± 27.2 491.8 ± 135.6 0.87 2.1 156SJGB3 11.10 ± 0.29 ments, various carbon isotopic trends for light hydrocarbon alkanes for Happo #1 sample. A new alternative model is 551.0 ± 25.8 530.2 ± 122.9 0.88 -0.9 157SJGB3 11.24 ± 0.36 0.0582 ± 0.0047 549.2 ± 16.9 546.8 ± 37.0 536.7 ± 185.4 0.65 -0.4 158SJGB3 products other than methane in some FTT experiments 0.0588 ± 0.0039 550.4 ± 14.8 552.3 ± 31.1 559.9 ± 150.8 0.84 0.3 159SJGB3 n-alkane products, which were sampled at relatively early 0.0602 ± 0.0044 558.1 ± 17.7 568.8 ± 35.5 611.8 ± 166.1 0.63 1.9 160SJGB3 hand, the 813C value of CO (starting material) fit the linear 0.0572 ± 0.0020 485.3 ± 14.5 487.8 ± 18.4 499.5 ± 79.0 0.60 0.5 161SJGB3 11.71 ± 0.39 0.0567 ± 0.0039 528.3 ± 16.8 519.2 ± 31.4 479.4 ± 160.2 0.74 -1.7 162SJGB3 10.96 ± 0.38 0.0547 ± 0.0039 562.9 ± 18.9 531.7 ± 33.7 400.1 ± 170.0 0.57 -5.5 LG1 Coarse-Grained leucogabbro 206SJLG1 11.59 ± 0.31 0.0594 ± 0.0046 533.6 ± 13.8 543.1 ± 35.1 583.1 ± 176.6 0.90 1.8 207SJLG1 11.62 ± 0.27 rocks in Canada (Sherwood Lollar et al., 2008). In those 532.1 ± 11.8 553.8 ± 28.1 644.2 ± 136.8 1.04 4.1 208SJLG1 11.56 ± 0.32 0.0551 ± 0.0037 534.7 ± 14.4 512.5 ± 30.2 414.8 ± 159.7 0.70 -4.2 209SJLG1 11.87 ± 0.34 0.0608 ± 0.0056 521.3 ± 14.5 542.6 ± 41.2 633.1 ± 211.5 0.67 4.1 210SJLG1 11.92 ± 0.32 0.0602 ± 0.0035 519.3 ± 13.5 536.5 ± 27.0 observed carbon isotopic distribution among all Ci—C5 0.74 3.3 211SJLG1 11.52 ± 0.31 ural hydrocarbons from Precambrian Shield crystalline 536.4 ± 13.9 570.1 ± 36.9 #1 sample. However, the expected isotopic relationship 0.67 6.3 216SJLG1 11.47 ± 0.33 conducted by using CO and H2 as the starting materials 538.9 ± 15.0 553.4 ± 39.6 613.5 ± 197.5 0.65 2.7 217SJLG1 11.20 ± 0.36 0.0531 ± 0.0061 551.4 ± 16.9 511.0 ± 48.5 the isotopic compositions of the lower-molecular weight 0.60 7.3 218SJLG1 in their model) to the growing hydrocarbon chain without 0.0608 ± 0.0041 530.2 ± 14.8 specificisotope compositions 631.2 ± 153.6 0.95 3.7 219SJLG1 11.65 ± 0.34 0.0602 ± 0.0046 531.0 ± 14.8 546.5 ± 34.8 611.7 ± 173.3 0.84 2.9 220SJLG1 11.57 ± 0.38 0.0558 ± 0.0053 534.3 ± 16.9 517.7 ± 41.8 445.0 ± 226.4 0.66 -3.1 221SJLG1 alyzed FTT reaction. As for the Sherwood Lollar et al. 0.0598 ± 0.0078 530.0 ± 18.3 542.5 ± 58.8 of the produced methane was approximately 10%o lower 0.53 2.4 226SJLG1 11.65 ± 0.35 0.0524 ± 0.0051 530.8 ± 15.3 489.8 ± 40.4 302.4 ± 240.3 0.81 7.7 227SJLG1 11.92 ± 0.36 0.0578 ± 0.0049 519.4 ± 15.0 519.7 ± 36.9 521.0 ± 195.8 0.73 0.1 228SJLG1 12.29 ± 0.42 0.0588 ± 0.0057 504.3 ± 16.5 514.3 ± 42.4 559.0 ± 226.8 0.74 2.0 229SJLG1 11.45 ± 0.29 0.0583 ± 0.0040 539.6 ± 13.0 540.0 ± 31.0 carbon number. This prediction is consistent with the 0.91 0.1 230SJLG1 compared to methane. In contrast, the 813C values of the addition of a carbon atom. One has been developed to molecular-weight alkanes (Des Marais et al., 1981; gas-phase radical recombination reactions in an electrical significant differences in isotopic trends among various 1.54 3.0 231SJLG1 9.33 ± 0.51 0.1521 ± 0.0207 656.2 ± 34.4 1196.0 ± 108.9 2370.0 ± 252.7 0.60 82.3 236SJLG1 11.70 ± 0.35 0.0543 ± 0.0054 528.6 ± 15.2 501.8 ± 42.0 381.5 ± 242.0 0.64 -5.1 237SJLG1 11.43 ± 0.34 0.0624 ± 0.0037 540.8 ± 15.6 kinetic fractionations during recombination due to 12CH3 et al., 1998; Fu et al., 2007; Taran et al., 2007; McCollom 0.95 5.4 238SJLG1 11.39 ± 0.34 0.0632 ± 0.0056 542.6 ± 15.6 formation via abiotic polymerization 716.3 ± 198.9 0.59 6.4 239SJLG1 11.53 ± 0.32 0.0614 ± 0.0056 536.1 ± 14.4 558.7 ± 42.1 higher hydrocarbons are predicted by mass balance from 0.57 4.2 240SJLG1 11.67 ± 0.32 0.0571 ± 0.0031 530.1 ± 13.8 523.2 ± 25.3 Telling et al., 2013) (Fig. 6c). The distribution of carbon 0.98 -1.3 241SJLG1 11.92 ± 0.32 0.0607 ± 0.0049 519.4 ± 13.5 540.2 ± 36.2 628.9 ± 184.8 0.84 4.0 (%) ol x [-(aenseedoz) / (ae nsez/doz)] se pauyap s! (% a) aouepop jo aaiap au 352 point of CH4, using only C2—Cs dataset (Fig. 5b): Oeyama body (Tsujimori et al., 2000), ca. 444-469 Ma from the Sm-Nd isotopic data of whole-rock, clinopyroxene and plagioclase Wakasa body (Nishimura and Shibata, 1989), ca. 430 Ma and ca. were generated and also processed the data based on this newly ob- 301-379 Ma from the Wakamiya-Sasaguri area in Northern Kyushu tained zircon age. (Nishimura and Shibata, 1989; Shibata et al., 1977). Fine-grained hornblende gabbro from Ashidachi body yielded ca.343Ma 3. Analytical procedures (Shibata et al.,1979). On the other hand, Nomo and Yamaga metagabboic bodies that are the western extension of Sangun- In this simple polymerization model, the observed car- Renge Belt and have arc/back-arc affinity yielded ca. 457-480 Ma sis was carried out using fused beads of the samples in RIGAKU ZSX- and ca. 477 Ma (Igi et al.,1979; Nishimura and Shibata, 1989; 101e system at Hiroshima University. The Rh-W dual anode tube was Shibata et al., 1977). Furthermore, recently, zircon U-Pb ages of ca. used to irradiate the sample. The analytical precision (1o) of major ele- 472-532 Ma from Osayama area (Tsujimori et al., 2005; Tsutsumi ments ranges between 0.001 and 0.14 wt%, while that for minor ele- et al, 2010) and ca. 497-519 Ma from Itoigawa-0mi area To evaluate the sensitivity of 813C-CH4 to the linear (Kunugiza et al., 2002; Tsutsumi et al., 2010) have been reported The electron back-scattered images (BSI) and cathodoluminescence from the jadeitite associated with serpentinite melange that under- (CL) images of the separated zircon grains were taken using the scan- lies the nappe of the Oeyama ophiolite. And these ages are ning electron microscope (SEM: JEOL JSM 750OF) at Hiroshima Univer- interpreted for the timing of their metasomatism during their sub- sity. These images are used to assess size, shape and internal structure of duction. The Oeyama ophiolite is thought to be older than these zir- the zircon grains, as well as identification of any mineral inclusions be- con U-Pb ages, possibly the oldest ophiolitic complex in the eastern tions; 813c# = 8l3c*-&p = (intercept) + (slope). The error bar of margin of the Asian continent, but no concrete dates were available The mineral phase chemical analysis was carried out by electron so far. In this study, zircon grains are found in two samples of the probe microanalyzer (EPMA: JEOL JXA 8200 Superprobe) at the Natural coarse-grained diallage gabbro in the Saijo body. This is the first suc- Science Center for Basic Research and Development (N-BARD), Hiro- cessful separation of zircon from the rocks of the Oeyama ophiolite. shima University. The operating conditions were 15 kV of accelerating We measured their U-Pb ages and chemistry using laser-ablation in- voltage, 20 nA beam current and 1-2 μm of probe diameter. Both natural ductively coupled mass spectrometer (LA-ICP MS). Furthermore, Calculation of the linear regression equation for a 8l3C data set of (a) 0.09 (b) 0.09 GB3 LG1 regression equation for a 8l3C data set of C2 to Cs. Gray area Note that this simple polymerization model presupposes 6 P 0.07 6 n 0.07 n-C4 20 Pb/ Pb/ 0.06 Although the r2 value decreases, the coeffcients of Eq. -60 640 -60 600 640 600 480 813Cn = 8.2 × (1/n) - 42.0 (r² = 0.75) 0.05 0.05 estimated intercept and slope of the regression line. 0.04 9.5 10.5 11.5 12.5 13.5 9.5 10.5 11.5 Based on the polymerization model derived in this 13.5 238 U/206pb data study, the ideal polymerization equation for Happo #1 errorsymbolsare2a (c) number ‘n’ includes one atom of C# and (n-1) atoms of GB3 (d) ± 3.3%o. The theoretical carbon isotopic composition of LG1 Fig. 5. Carbon isotope values plotted against the inverse carbon 580 580 with a coefcient of determination (r²) of 0.97. The &p, value for Happo #1 sample can only be explained by providing with each other to form a longer carbon chain (e.g., pentane ge ige 560 560 VPDB a 0.6 238 0.4 540 540 P 206 206 520 520 500 500 Fig. 7. Tera-Wasserburg concordia diagrams of (a) GB3 and (b) LG1 and Weighted average age diagrams of 206pb/238U ages of (c) GB3 and (d) LG1. (a-b) Black ellipses represent concordant data, gray ellipses represent rejected data, and gray dashed ellipses represent discordant data. Table 4 LA-ICP-MS REE and trace element data from Saijo mafic-ultramafic body. 203SJGB3 204SJGB3 205SJGB3 206SJGB3 209SJGB3 211SJGB3 212SJGB3 215SJGB3 216SJGB3 218SJGB3 221SJGB3 222SJGB3 223SJGB3 224SJGB3 225SJGB3 303SJLG1 Y 2471.7 3150.2 2588.0 2408.1 2507.4 3096.6 3670.4 3355.4 3936.1 3620.8 3101.5 4183.9 2129.5 2555.3 2739.7 2369.3 Nb 11.34 5.80 8.58 8.00 12.23 5.39 5.53 6.94 15.87 5.25 10.81 16.65 8.26 7.96 8.05 6.15 La <LLD <LLD 0.0081 0.0200 <LLD 0.0172 0.0462 <LLD <LLD 0.0024 <LLD 0.0922 0.0667 <LLD <LLD <LLD Ce 25.56 26.82 17.66 22.24 24.96 21.82 24.64 22.13 31.14 28.85 29.67 35.36 20.22 21.59 20.01 18.20 Pr 0.034 0.037 0.116 0.055 0.043 0.062 0.233 0.150 0.202 0.052 0.288 0.179 0.049 0.051 0.081 0.079 PN 1.678 2.501 2.050 2.479 1.110 3.394 4.104 6.362 3.589 2.612 4.416 4.777 1.804 2.486 1.775 2.151 Sm 5.293 7.704 6.215 4.287 4.094 10.082 11.664 13.210 8.965 9.543 6686 10.943 4.524 7.799 4.466 4.442 Eu 0.815 1.109 1.128 0.950 0.658 1.499 1.903 2.003 1.786 1.669 1.533 2.006 0.869 0.914 0.937 1.291 Gd 34.96 55.96 38.69 37.41 33.81 63.06 65.09 65.84 66.70 69.46 56.21 76.80 33.73 46.35 42.38 39.48 Tb 15.41 22.11 16.97 14.45 14.15 23.78 26.55 23.98 26.53 27.85 22.26 30.44 12.61 18.91 17.19 16.12 Dy 212.5 263.2 221.6 191.9 188.3 308.2 326.1 312.4 385.2 356.3 274.3 414.0 182.7 218.3 230.9 211.4 Ho 84.69 113.50 89.02 81.98 71.60 110.74 131.84 114.57 142.03 135.29 112.28 162.34 67.06 94.96 89.98 89.59 Er 415.3 508.1 429.0 377.2 383.3 527.9 626.8 539.5 676.7 570.0 514.7 770.6 326.3 401.8 408.5 390.1 Tm 94.94 109.72 95.71 85.47 91.35 120.20 124.31 111.35 149.26 130.92 118.18 172.25 83.09 108.54 86.27 88.17 Yb 933.2 1047.2 925.0 883.2 840.2 1068.3 1190.7 1167.3 1673.3 1400.4 1257.1 1658.2 740.6 1039.4 879.3 956.9 Lu 165.4 183.6 160.2 155.6 158.6 189.2 199.1 180.6 264.0 203.0 197.2 265.5 125.7 166.5 139.5 153.3 Hf 15,066.9 13,608.4 11,965.0 13,319.2 14,849.0 13,811.9 13,811.9 12,970.7 11,388.4 14,480.1 12,732.4 12,732.4 13,818.7 14,055.2 13,320.0 13,209.0 Pb 24.16 40.47 13.17 18.27 26.18 19.47 19.17 23.53 24.21 24.80 28.71 29.46 19.87 26.28 18.91 14.72 Th 47.11 113.10 26.62 36.76 45.09 46.96 52.18 66.25 49.79 61.34 79.01 62.92 28.15 61.64 43.82 31.64 U 81.68 148.92 43.35 56.80 78.62 65.38 65.56 76.85 84.11 90.37 85.31 84.08 58.37 81.72 56.75 53.60 Ce/Ce* 141.3 164.4 163.8 58.2 633.2 67.5 86.7 Eu/Eu* 0.18 0.16 0.22 0.23 0.17 0.18 0.21 0.21 0.22 0.20 0.20 0.21 0.22 0.15 0.21 0.30 304SJLG1 305SJLG1 306SJLG1 309SJLG1 310SJLG1 311SJLG1 312SJLG1 315SJLG1 316SJLG1 317SJLG1 318SJLG1 321SJLG1 322SJLG1 323SJLG1 324SJLG1 327SJLG1 328SJLG1 329SJLG1 1667.5 2056.9 2404.9 2567.9 2076.9 2349.1 1834.4 2190.3 3101.4 2348.0 2203.2 2992.8 2719.9 2563.3 1388.1 2649.0 2435.2 1655.3 -343 7.00 7.04 9.73 6.84 7.00 7.33 5.65 7.39 5.98 8.47 6.98 7.61 8.46 13.54 4.95 7.70 5.78 5.52 <LLD <LLD 0.0098 0.0441 <LLD <LLD <LLD <LLD 0.0756 0.0552 <LLD <LLD 0.0592 0.0238 <LLD 0.1656 0.0452 0.0606 (2019) 13.66 13.86 24.32 21.77 17.33 17.45 11.52 19.75 17.19 22.77 19.37 23.24 15.78 25.78 12.38 21.47 16.43 12.34 0.097 0.021 0.029 0.119 0.132 0.024 0.008 0.015 0.392 0.057 0.042 0.111 0.171 0.066 <LLD 0.170 0.188 <LLD 1.250 1.671 1.006 3.340 1.375 1.668 2.360 2.276 5.143 2.940 1.905 2.181 2.432 1.138 0.907 2.418 1.559 1.713 3.442 4.652 5.879 7.360 4.706 5.876 4.500 5.318 11.439 4.673 5.088 7.103 8.514 5.413 3.088 5.806 5.495 3.344 360 0.921 0.554 0.852 2.063 1.035 1.109 0.792 1.371 2.569 1.447 1.297 1.431 1.396 0.800 0.700 1.591 1.342 0.705 27.24 30.97 34.11 53.55 33.14 33.47 26.36 32.14 57.23 41.46 36.79 47.66 41.94 38.47 18.97 41.86 35.63 25.29 11.64 13.28 15.62 19.77 14.15 14.77 12.91 14.14 24.11 16.40 15.47 20.31 19.50 16.10 9.20 18.03 16.00 10.28 150.2 178.0 213.1 279.5 189.9 196.1 159.2 189.4 281.5 227.3 225.0 266.1 248.3 212.2 117.3 233.9 205.3 141.2 62.92 71.66 82.54 104.71 68.70 72.95 59.50 78.19 107.82 85.63 78.61 103.00 102.00 88.50 45.51 89.52 76.82 52.09 325.8 325.8 430.5 501.0 329.8 382.9 282.6 359.5 506.5 381.4 352.1 479.3 482.6 386.3 220.1 368.9 355.1 272.6 67.00 77.82 103.67 100.80 84.01 80.42 69.15 77.83 102.41 89.31 88.34 115.29 102.92 100.78 54.76 93.95 83.43 64.26 692.0 899.9 954.0 731.0 739.8 949.7 661.7 757.5 883.1 927.0 784.1 1011.0 1056.7 973.5 551.1 985.4 860.4 647.7 112.7 135.0 166.7 167.5 135.0 144.5 110.8 140.7 170.2 145.0 131.1 170.0 180.2 167.2 92.0 158.3 150.8 100.8 13,292.1 12,376.2 14,854.9 11,879.3 11,879.3 12,738.3 13,098.7 13,197.1 11,357.0 12,997.8 12,779.9 12,446.6 12,478.9 14,473.3 13,756.8 12,633.2 12,556.9 13,086.0 13.97 16.87 25.81 18.39 17.68 15.70 12.03 27.09 11.31 16.71 16.76 24.47 18.23 26.67 9.46 17.66 11.97 14.81 21.18 34.51 48.77 56.35 29.81 32.27 26.80 72.55 37.65 29.24 30.34 55.85 30.76 43.62 14.22 31.69 25.20 25.93 59.25 58.26 44.67 56.13 86.46 42.39 84.33 50.72 41.51 88.12 49.00 53.00 83.41 56.74 28.04 49.36 40.20 42.76 353.7 73.7 24.5 99.5 38.5 159.5 31.4 43.7 0.29 0.14 0.18 0.32 0.25 0.24 0.22 0.32 0.31 0.32 0.29 0.24 0.23 0.17 0.28 0.31 0.29 0.23 CecN/CecN* = Cecn/(LacnPrcn)1/2. Euce/EucN*= Euce/(SmcNGdcN)/2 LLD: lower limit of detection. 3 354 K.Kimura,Y.Hayasaka/Lithos 342-343(2019) 345-360 100000 standard sample) during this study were 0.71025 ± 0.00002 (20, n = 6) and 0.512709 ± 0.000014 (20, n = 6) respectively. 10000 GB3 0.512782 for JB-1a. We estimate an error of 0.5% for the 1000 × 10-12/y and 1.42 × 10-11/y, respectively (Lugmair and Marti, 1978; Steiger and Jager, 1977). 100 C1 10 4. Results 4.1.Sample descriptions 0.1 LG1 4.1.1. Coarse-grained diallage gabbro (GB3, 92061305, 5-19-4, 5-20-3, 5-20-9,5-25-10C, GB2 and GB4) The first two samples, i.e. GB3 (Fig. 3a) and 92061305 belong to the 0.01 LaCePrNd SmEuGdTbDyHoErTmYbLu Unit Ill of the Saijo body (Fig. 2c). The rest of the six samples belong to the Unit Ill of the Ashidachi body. The sample locations are mentioned Fig. 8. C1 chondrite-normalized REE pattern of the analyzed zircon in GB3 and LG1. on the geological map (Fig. 2b, c). The main constituent minerals of these samples are plagioclase and clinopyroxene (diallage) with For LA-ICPMS, zircon U-Pb isotope analysis was performed using minor amounts of orthopyroxene and ilmenite (Fig. 4a, b). Relatively 213 nm Nd-YAG Laser (New Wave Research UP-213) coupled with fresh plagioclase shows lamellar twining, and some plagioclase grains Agilent 7500 ICP-MS at the Department of Earth and Planetary Systems are found to be saussuritized. Clinopyroxene occurs as a diallage that Science, Hiroshima University. The detailed analytical methodology is has well-developed cleavage. Orthopyroxene is completely altered described by Katsube et al. (2012),their Table 1. The apparatus contains to bastite. Secondary minerals are mostly chlorite, actinolite and a mixed He-N2-Ar carrier gas system equipped with small volume abla- magnesiohornblende forming at the grain boundary and along the tion cell, sample aerosol stabilizer (buffering chamber) and charcoal fil- cleavage planes of diallage. GB3 and 92061305 collected from Saijo ter attachment. The used spot size of the laser was 25 μm. Zircon grains are measured in a mixed sequential order to avoid the influences of 100 time-depending changes in the instrumental condition. Raw data (a) 口GB3 were processed using the data reduction program PepiAGE (Dunkl LG1 et al., 2008). The processed data of both analytical procedures have Continental zircon ·Yakuno ophiolite been finally used for statistical plotting using Isoplot/Ex (Version 3; 10 Ludwig, 2003). The isotopic ratios and age data are quoted at the esti- mation error of 2o level, whilst, the weighted mean is given at the 95% confidence level. The obtained data points are characterized on the basis of their U content and Th/U ratio. In each sample, the weighted average is calculated and the probability density plots are drawn with data having the near-concordant age values within their estimation Ocean crust zircon error. Zircon standard of FC1 (206pb/238U age of 1099.0 ± 0.6 Ma; 0.1 Paces and Miller, 1993) was used for correction of U-Pb ratio, and glass standard of NIST SRM 610 was used for correction of Th/U ratio, and zircon grains of KO1 (TIMS 206pb/238U age of 95.6 Ma; Herzig 0.01 et al., 1998) was used as consistency standard. During this analytical 5000 10000 15000 20000 25000 30000 35000 session, the weighted mean 206pb/238U ages of K01 were 96.4 ± Hf (ppm) 2.0 Ma and 93.9 ± 1.3 Ma. 100 Before the analysis of the trace elements in zircon using LA-ICP (b) 口GB3 MS, Hf concentration was measured using the EPMA (JEOL JXA LG1 8200 Superprobe) data produced at the Natural Science Center for ·Yakunoophiolite Basic Research and Development, Hiroshima University in order to 10 Continental zircon normalize the trace element concentrations. The operating condi- tions of EPMA were 15 kV accelerating voltage, 20 nA beam current and 6 μum beam diameter. The used spot size of the laser was 15 μum in diameter. For the Sm-Nd and Rb-Sr isotope analysis, plagioclase and clinopyroxene were separated using an isodynamic separator and heavy liquid, then leached by diluted hydrochloric acid. The deter- 0.1 mination of isotope ratios of 87sr/86Rb and 143Nd/144Nd and elemen- tal abundances of Rb, Sr, Nd and Sm by isotope dilution method was Oceancrustzircor carried out using MAT261-type (modified from MAT260) thermal 0.01 10 100 1000 10000 100000 Institute for Studies of Earth's Interior, Okayama University. Detail Y (ppm) of the analytical procedure is described in Kagami et al. (1992). 87sr/86sr ratios were normalized to 86sr/88sr = 0.1194 and Fig. 9. (a) Hf vs. U/Yb and (b) Y vs. U/Yb diagrams of the studied zircon grains from both coarse-grained diallage gabbro and leucogabbro (after Grimes et al, 2007). Yakuno 87Sr/86Sr ratio for NBS987 and 143Nd/144Nd ratio for JB-1a (GSJ ophiolite derived from island arc and/or back-arc crust is plotted in the mixed zone whereas, the presently studied samples are all plotted in the oceanic crust region. K.Kimura,Y.Hayasaka/Lithos 342-343(2019)345-360 355 body are relatively rich in metamorphic amphibole and opaque min- 46 and 49 wt%, while that from Saijo body falls in a narrow range from erals than those from the Ashidachi body. Anorthite content of plagio- 51 to 52 wt%. Magnesium content (Mg#) of gabbroic samples from clase ranges from 60 to 65%, and pyroxene chemistry is close to that of Ashidachi body shows slightly higher value from 63 to 74% than that diposide (Wo44-47 En44-48 Fss-11) (Table 1). Zircon grains could be sep- from the Saijo body (47 to 63%). Immobile HFSEs show values below 1 arated from only the sample GB3. in the MORB-normalized spider diagram (Fig. 5b). 4.1.2. Coarse-grained leucogabbro (LG1) 4.3. Zircon U-Pb age and chemistry This sample is collected from the same outcrop of GB3 that belongs to the Unit Il of the Saijo body. LG1 has direct contact with GB3, Among all the studied samples zircon grains could be successfully which is characterized by the chilled margin (Fig. 3b, c). LG1 contains separated and measured only from the samples GB3 (coarse-grained major minerals of plagioclase (60-80% by volume) and altered diallage diallage gabbro) and LG1 (coarse-grained leucogabbro). (20-35% by volume) with a minor amount of ilmenite and zircon Zircon grains from the sample GB3 are 50-200 μm in size, most of (Fig. 4c, d). Occasionally, titanite and epidote-rich veins are present. them are <100 μum. On the other hand, zircon grains from the sample Secondary minerals are mostly chlorite, actinolite and magnesiohor- LG1 are larger in size (100-200 μm, the majority are larger than 150 nblende changed from diallage. The modal abundance of zircon is μm). Most of the grains show the euhedral shape and oscillatory zoning higher than that of the GB3. Anorthite content of plagioclase (56 to in the SEM-CL images especially for large grains (Fig. 6a, b). ( ) n no n n q s ( We have measured 44 spots from 28 zircon grains separated from GB3 (Table 3). Analyzed data of 40 points out of 44 lie on the concordia 4.1.3. Fine-grained hornblende gabbro (162, 171, 5-25-9B, 6-13-4) line within their analytical error, which have discordance <10% (Fig. 7a). Except one sample from the Saijo body (6-13-4), all other the stud- ied samples belong to the Unit Il of the Ashidachi body. The main con- except for statistically rejected 3 data points yields an age of 545.4 ± stituent minerals of these samples are plagioclase and hornblende. 2.6 Ma (2o, MSWD = 1.4) (Fig. 7c). Age probability density diagram Clinopyroxene rarely occupies the core part of hornblende. It is difficult of the concordant data shows a nearly symmetric single peak of to pinpoint whether the hornblende is igneous or metamorphic in so eneed origin. A minor amount of quartz is also present. Plagioclase from fine- 18 zircon grains separated from LG1 have been analyzed (Table 3). A grained hornblende gabbro is generally fresh than those in the coarse- total of 24 spots from these grains for LG1 yield 23 concordant data grained diallage gabbro. Accessory minerals are mostly apatite, zircon pm pe is on tos wy a se hego (ae ) and opaque minerals. Secondary minerals are chlorite, calcite and actin- olite. Pyroxene chemistry is close to that of augite to endiopside 532.4 ± 3.1 Ma (20, MSWD = 1.20) (Fig. 7d). The 206pb/238U age prob- (W036-45 En38-51 FS6-17) (Table 1). ability density diagram shows a nearly symmetric single peak at ca. 534 Ma. Th/U ratio is ranging between 0.53 and 1.04, except for one 4.2. Bulk chemical compositions point with a higher value of 1.54. Trace element concentrations of 15 points from 10 zircon grains of Chemical compositions of the analyzed 9 samples are shown in the sample GB3 and 19 points from 11 zircon grains separated from Table 2. Gabbroic samples from Ashidachi body contain SiO2 in between the sample LG1 are also measured using LA-ICP MS (Table 4). Fig. 8 Table 5 Sm-Nd and Rb-Sr isotope composition of coarse-grained diallage gabbro and fine-grained hornblende gabbro from Ashidachi and Saijo bodies. Sm (ppm) Nd (ppm) 147Sm/144Nd PNt+/PNet1 Rb (ppm) Sr (ppm) 87Rb/86Sr 87sr/86sr Sample name type Coarse-grained diallage gabbro Ashidachi body 5-19-4 WR 0.5274 1.0436 0.305622 0.513417 ± 0.000007 3.832 161.52 0.068603 0.702862 ± 0.000011 cpx 1.1488 2.0928 0.332011 0.513510 ± 0.000013 pl 0.0963 0.3738 0.155744 0.512912 ± 0.000026 5-20-3 WR 0.6080 1.5365 0.239319 0.513184 ± 0.000013 6.188 210.57 0.084995 0.704108 ± 0.000011 cpx 2.0043 4.2271 0.286767 0.513347 ± 0.000009 pl 0.2502 1.0274 0.147271 0.512790 ± 0.000021 5-20-9 WR 0.5525 1.2321 0.271165 0.513278 ± 0.000008 6.023 286.90 0.060718 0.703571 ± 0.000012 cpx 1.2096 2.4000 0.304809 0.513406 ± 0.000009 pl 0.1108 0.4773 0.140398 0.512776 ± 0.000028 5-25-10C WR 0.8650 2.0732 0.252321 0.513206 ± 0.000009 10.754 162.72 0.191169 0.704772 ± 0.000017 xd 1.5432 3.3864 0.275601 0.513329 ± 0.000023 1.497 11.15 0.388338 0.703672 ± 0.000028 pl 0.2763 1.0003 0.167039 0.512865 ± 0.000021 Saijo body 92,061,305 WR 0.9878 2.3920 0.249744 0.513209 ± 0.000009 35.603 251.57 0.409296 0.703518 ± 0.000009 xd 3.4747 7.2468 0.289992 0.513369 ± 0.000010 pl 0.2694 1.2634 0.128967 0.512807 ± 0.000014 Fine-grained hornblende gabbro Ashidachi body 5-25-9B WR 2.9731 8.6036 0.208994 0.512957 ± 0.000009 6.023 286.90 0.060718 0.703571 ± 0.000012 xd 4.7465 13.0745 0.219546 0.512978 ± 0.000007 pl 0.2624 1.1203 0.141620 0.512811 ± 0.000014 162 WR 3.7244 11.6011 0.194101 0.512990 ± 0.000009 8.970 210.81 0.12307 0.703995 ± 0.000017 171 WR 2.0353 6.2858 0.195772 0.512959 ± 0.000012 7.419 305.05 0.07034 0.704076 ± 0.000013 Saijo body 6-13-4 WR 3.6722 10.7439 0.206698 0.512968 ± 0.000009 18.450 128.74 0.414532 0.704929 ± 0.000011 cpx 9.0862 30.9002 0.177823 0.512879 ± 0.000011 0.835 412.69 0.005853 0.704365 ± 0.000008 pl 0.6241 2.0962 0.180046 0.512892 ± 0.000009 20.532 175.12 0.339116 356 K.Kimura,Y.Hayasaka/Lithos 342-343(2019) 345-360 shows the C1 chondrite-normalized REE diagram. Both the samples of ± 500 Ma (20, MSWD = 9.6) for the sample 5-20-3,584 ± 27 Ma GB3 and LG1 show REE pattern with HREE enrichment, strong positive (20, MSWD = 0.017) for the sample 5-20-9, 629 ± 530 Ma (20, Ce anomaly (Ce/Ce* = 25-164 except two points showing higher values MSWD = 4.6) for the sample 5-25-10C, and 529 ± 310 Ma (20, of 354 and 633) and negative Eu anomaly (Eu/Eu* = 0.14-0.32). How- MSWD = 8.4) for the sample 92061305 (Fig. 10a). Some of these data ever, overall REE values of the sample GB3 is slightly higher than those yielded high error values due to large error in plagioclase isotopic ratios. of LG1. Moreover, discrimination diagrams like Hf vs. U/Yb diagram and However, Sm-Nd whole rock isochron age from the five whole rock data Y vs. U/Yb diagram are known to differentiate effectively between the is 566 ± 95 Ma (2o, MSWD = 4.3). The initial Nd ratio for the five sam- zircon grains formed in the oceanic crust and that formed in the conti- ples ranges between 0.51217 and 0.512383 and the re-calculated values nental crust (Grimes et al., 2007). Both the above mentioned samples for those five samples adapting whole-rock Nd ratio and zircon age of are plotted in the field of ocean crust zircon region (Fig. 9a, b). and etSr are also calculated using the zircon 206pb/238U age of 545 Ma 4.4. Nd and Sr isotope systematics (as estimated from the analyzed coarse-grained diallage gabbro, being the typical component of the oceanic crust). The e'Nd and &'Sr are vary- 4.4.1. Coarse-grained diallage gabbro ing in the range from +7.2 to + 7.7 and from -50 to -6, respectively. In Sm-Nd and Rb-Sr isotopes are measured for whole-rock samples of eNd vs. 8*Sr isotope correlation diagram (Fig. 10c), the studied samples 92061305, 5-19-4, 5-20-3, 5-20-9 and 5-25-10C. For Sm-Nd isotopic plot in the field of positive e'Nd and negative &'Sr. measurements, separated clinopyroxene and plagioclase are measured, 4.4.2.Fine-grained hornblende gabbro also measured. The analyzed data are listed in Table 5. From the five Sm-Nd isotopes are measured for whole-rock samples of 5-25-9B, 6- Sm-Nd whole rock-mineral internal isochron calculations yield the 13-4, 162 and 171. Clinopyroxene and plagioclase are separated and ages of 517 ± 25 Ma (20, MSWD = 0.20) for the sample 5-19-4, 602 measured for samples 5-25-9B and 6-13-4. The analyzed data are listed data-point error crosses are 20 0.5140 data-point 0.5140 (a) (b) 0.5136 0.5136 5-19-4 92061305 5-20-9 5-20-3 5-19-4 5-20-9 144 5-25-10C 92061305 0.5128 43 0.5124 0.5124 5-25-10C 5-20-3 0.5120 0.5120 0.0 0.1 0.2 0.3 0.4 0.0 0.1 0.2 0.3 0.4 10 0.5140 (c) (d) DM Seawater-rockinteraction 8 0.5136 6 EPRyoung St. Helena 6-13-4 PN+ Azores 0.5132 162 HIMU French Polynesian 144 d 0.5128 171 5-25-9B C-gabbro WR EM-1 0.5124 9- t=545 Ma -60 -40 -20 20 0.5120 εSr(t) 0.0 0.1 0.2 0.3 0.4 Fig. 10. (a) 147sm/144Nd vs. 143Nd/144Nd diagram of coarse-grained diallage gabbro.Reference isochron of 545 Ma is used to recalculate the initial Nd values. Sample number 92061305 is from the Saijo body, whereas all the data are from the Ashidachi body. (b) &Srvs.&Nd plot of coarse grained diallage gabbro samples.The time integration was done using thevalue of 545 Ma as derived from the zircon U-Pb age estimation. The initial epsilon Nd and Sr (=e'Nd and e'Sr) were calculated from the CHUR values of 147Sm/144Nd = 0.1967,143Nd/144Nd = 0.512638, 87Rb/86Sr = 0.0839 and 87Sr/86Sr = 0.7045, respectively. (c) 147Sm/144Nd vs. 143Nd/144Nd diagram of fine-grained hornblende gabbro. Reference isochron of 340 Ma is used to recalculate the initial Nd values. Sample number 6-13-4 is from the Saijo body, whereas all the data are from the Ashidachi body. K.Kimura, Y.Hayasaka / Lithos 342-343 (2019) 345-360 357 in Table 5. From two Sm-Nd whole rock-mineral internal isochron plots from the whole rock chemistry. However, the whole rock e'Nd of the yield the ages of 328 ± 30 Ma (2o, MSWD = 0.080) for the sample 5- present samples range from +7.2 to +7.7 and having an isotopic com- 20-9B, and 452 ± 65 Ma (2o, MSWD = 0.76) for the sample 6-13-4. position similar to MoRB (Hofmann, 2o07).Furthermore, zircon trace And the isochron age for all data plotted from the four samples is 343 element composition of the studied rocks indicates oceanic crust origin ± 82 Ma (2o, MSWD = 6.5) (Fig. 10d). The initial Nd ratio for the five (Fig. 9). Oceanic crust contains not only MORB but also arc or back-arc samples ranges between 0.512169 and 0.512382. basin basalt. Trace element composition of zircon from the typical arc and back-arc origin, e.g. Yakuno ophiolite (our unpublished data) oc- 5. Discussion cupy the transition field in Hf vs. U/Yb and Y vs. U/Yb diagrams (Fig. 9). However, the present data exclusively plot in the oceanic 5.1. Age and tectonic affinity crust domain. This leads us to consider the depleted mantle origin oce- anic crust as the parentage for the coarse-grained gabbro of the Oeyama In terms of the structural position just above the ultramafic rocks ophiolite. and a layered structure in some outcrops, all of the presently studied Zircon grains, being euhedral in shape with oscillatory zoning and samples may correspond to the mafic cumulates of the ancient oceanic Th/U ratio ranging between 0.5 and 1.0 indicate that the age values crust. However, the corresponding dolerite-basalt sequence has not measured from each sample represent their respective age of magmatic been found in the Oeyama ophiolite yet. Therefore, it is diffcult to dis- crystallization. The five-point whole rock Sm-Nd isochron age of 566 ± cuss the nature of the source magma of the Oeyama ophiolite only 95 Ma is comparable with zircon 206pb/238U age values of 545.4 ± (a) latest Neoproterozoic Oeyama ophiolite Unit I (b) early Cambrian Oeyama Unit I Jadeitite Nomo-Yamaga (c) Ordovician arc/back-arc Oeyama Unit I Oeyama Unit Il (d) early Carboniferous Sangun-Renge N-Y arc Unit I metamorphic rocks OeyamaUnit IlI Fig.11. Schematic diagrams of progressive development of the oceanic crust (a) scenario at ca.545 Ma forming the early oceanic crust at MOR of Panthalassa after the breakup of the of cumulates progressed tillca. 400 Ma ( d) the switching of the subduction polarity at ca. 400 Ma and final obduction to the accretion system and related metamorphism up to <300 Ma. 358 K.Kimura,Y.Hayasaka/Lithos 342-343(2019)345-360 2.6 Ma and 532.4 ± 3.1 Ma within their error limit. This is the first report earliest subducted oceanic crust at the East Asia (Fig. 11b). The reported of Late Neoproterozoic zircon U-Pb age from the gabbroic rocks origi- age of ca. 530 Ma of jadeitite formation might imply the initiation of this nated from the depleted mantle in East Asia. Considering the oceanward subduction, when the Oeyama ophiolite Unit I was present paleocontinental distribution, this age postdates the rifting of the as a part of the overriding oceanic plate (Fig. 11b). A similar oceanward Rodinia assembly (800-700 Ma: Meert and Torsvik, 2003; Bose et al., subduction scenario is also proposed for the Ordovician Miyamori- 2016) and represents the post-Rodinia oceanic crust formation with Hayachine ophiolite, Northeast Japan (Ozawa et al., 2015). During the N-MORB affinity, and preserved as a remnant of the paleo-Pacific above-mentioned subduction, metasomatism of the Oeyama ophiolite crust. Moreover, the studied leucogabbro might have been derived (Machi and Ishiwatari, 2010; Nozaka, 2014), formation of the Ordovi- from the off-ridge magmatism from the view point of its younger age cian island arc (Igi et al., 1979; Nishimura and Shibata, 1989) and the and occurrence. On the other hand, fine-grained hornblende gabbro Ordovician to Devonian metamorphism of troctolitic to anorthositic cu- samples yielded an age of 328 and 452 Ma. Leaving aside the sample mulate of the Unit II (Nishimura and Shibata, 1989; Tsujimori, 1999; 6-13-4 that yielded the older age of 452 Ma, all data are plotted on Tsujimori et al., 2000) were occurred (Fig. 11c). Subsequently, the sub- 340 Ma referential isochron that has initial Nd values calculated by the duction direction was switched from oceanward to the present-day average of re-calculated initial Nd values for all samples (Fig. 10d). At continent-ward configuration at ca. 4o0 Ma forming the new accretion- this moment, the meaning of the 452 Ma age of one of the fine- ary complex that was the protolith of the Sangun-Renge metamorphic grained hornblende gabbro sample is unclear. Therefore, the protolith rocks. During this ongoing continentward subduction, the dyke of the of fine-grained hornblende gabbro is considered to be of ca. 340 Ma age. fine-grained hornblende gabbro and the basalt of the Unit Ill were The presence of amphibole in diallage gabbros, especially in the Saijo formed at the back-arc basin in the Carboniferous period (Hayasaka body, and the whole-rock geochemistry indicate that the Oeyama et al., 1995; Shibata et al., 1979) (Fig. 11d). This age roughly coincides ophiolite had suffered high temperature hydration. In addition, the with the age of metamorphism of the Sangun-Renge schist that under- Oeyama ophiolite suffered metasomatism at the forearc setting on the lies the Oeyama ophiolite (Ishiwatari, 1991a). mantle wedge (Machi and Ishiwatari, 2010; Nozaka, 2014). We con- In the circum-Pacific region, there are reports of some other latest sider that the formation of amphibole in the sample GB3 occurred dur- Neoproterozoic to early Cambrian ophiolites. These include Trinity ing the same process in this setting. However, the timing of the ophiolite (ca. 556-579 Ma） in Klamath Mountains at western metasomatism cannot be deciphered from the present data set. North America (Wallin, 1990; Wallin et al., 1988; Wallin et al., 1995), Ust'-Belaya ophiolite (ca. 542-575 Ma) at Far East Russia (Hayasaka et al., 2010; Ledneva et al., 2012; Palandzhyan, 2015; 5.1.1. Reassessment of the existing geological classification of the Oeyama Tikhomirov, 2010), Marlborough ophiolite (562 ± 22 Ma and 530 ophiolite ± 6 Ma) at New England Fold Belt of eastern Australia (Aitchison So far, the classification of different rock units of the Oeyama ophio- et al., 1992; Bruce et al., 2000). These ophiolites are estimated vari- lite was based on the rock type considerations (Kurokawa, 1985). ous tectonic setting such as slow-spread rifting basin, oceanic However, the coarse-grained diallage gabbro (ca. 545 Ma) shows within-plate setting, sea-floor spreading center, arc and back-arc en- interfingering relationship with the ultramafic rocks of Unit I of vironment (Bruce et al., 2000; Lindsley-Griffin, 1994; Mankinen Kurokawa ( 1985). Moreover, the fine-grained hornblende gabbro is in- et al., 2002; Palandzhyan, 2015; Yang and Seccombe, 1997). Though truded both coarse-grained diallage gabbro (ca. 340 Ma) and ultramafic the breakup of Rodinia is thought to occur during ca. 800-700 Ma rocks of Unit I of Kurokawa (1985). Their intrusive relationships in the (Meert and Torsvik, 2003), any ophiolite body older than 600 Ma field occurrence and the age data from the present study lead us to con- has not been found in the circum-Pacific region. This may indicate sider that the coarse-grained diallage gabbro can be reassigned to the Unit I of Kurokawa (1985). As per our refined classification, the Oeyama that the old initial stage oceanic crust was completely subducted into the mantle during the period of oceanward subduction in the ophiolite may be further grouped as, Unit 1a consists of the ultramafic same way as demonstrated by the tectonic models of Spaggiari rocks, unit 1b contains the coarse-grained diallege gabbro, unit 2 is et al. (2003) and Ozawa et al. (2015). After about 100-150 myr epidote-amphibolite grade ultramafic to mafic metamorphosed cumu- later, oceanic crustal subduction and arc back-arc system had been lates, unit 3a is fine-grained hornblende gabbro, and 3b is metabasalt as- initiated to be formed at different places along the paleo-Pacific mar- sociated with the fine-grained gabbro. gin, including that of the presently studied Oeyama ophiolite. 5.1.2. Tectonic evolution 6. Conclusions The radiometric age data previously reported and those from the present study on the Oeyama ophiolite and its associated rocks lead to · Zircon 206pb/238U age of 545.4 ± 2.6 Ma and 532.4 ± 3.1 Ma were ob- an evolution process that is summarized in the following section. Fig. tained from coarse-grained diallage gabbro and leucogabbro of the 11 depicts the schematic evolution of the oceanic crust until its final ac- Oeyama ophiolite (Saijo body). The whole rock Sm-Nd isochron age cretion to the Japanese Islands. of 566 ± 95 Ma, though with slightly larger uncertainty this age is The unit I (la and Ib) of the Oeyama ophiolite was a fragment of the oceanic lithosphere originated from the depleted mantle formed at the latest Neoproterozoic to early Cambrian zircon U-Pb age from gab- 545 Ma, i.e. that is 100-150 myrs after the dispersion of the continental broic rocks originated from the depleted mantle in the East Asia. blocks of the Rodinia supercontinent (Fig. 11a). The remnant of the typ- · The Nd isotopic composition and zircon chemistry indicate their N- ical oceanic lithosphere formed during this 100-150 myrs time span has MORB affinity. Coarse-grained diallage gabbro can be reassigned to not been found anywhere in the circum-Pacific region. This may indi- the Unit I based on its field occurrence, age and chemistry, whilst it cate that substantial amount of the oceanic lithosphere was recycled was earlier classified as a member of the Unit Ill by Kurokawa (1985). into the mantle. The absence of any ophiolite sequence in the Circum- · Oeyama ophiolite represents the oceanic crust formed from the de- Pacific belt in the time gap 100-150 myrs between the initiation of pleted mantle rocks postdating at least 150 myr after the reported Rodinia breakup and the presently studied ca. 545 Ma Oeyama ophiolite Rodinia breakup. During the subsequent oceanic plate subduction, possibly signified a time period of oceanward subduction, which caused the Oeyama ophiolite and its associated rocks represent a part of the to the overall consumption of the oceanic lithosphere connected to the latest Neoproterozoic obducted oceanic crust and Cambrian to Car- continental margin.Cambro-Ordovician jadeitite (Tsujimori et al., 2005; boniferous arc-back arc to subduction zone material. Tsutsumi et al., 2010) can be ascribed to one of the evidence of the K.Kimura,Y.Hayasaka/Lithos 342-343(2019)345-360 359 Acknowledgements Ishiwatari, A., Hayasaka, Y., 1992. Ophiolite nappes and blueschist of the inner zone of Southwest Japan. In: Kato,H, Noro, H. (Eds.), 29th Field Trip Guide Book 5. Geological Survey of Japan, pp.285-325. We thank Y. Shibata for his help with EPMA analysis; H. Kagami Ishiwatari, A., Tsujimori, T., 2003. Paleozoic ophiolites and blueschists in Japan and with the Sr-Nd isotopic analysis; and Y. Takahashi with the zircon Russian Primorye in the tectonic framework of East Asia: a synthesis. Island Arc 12, U-Pb dating and zircon trace element measurement. Critical com- 190-206. Ishiwatari, A., Sokolov, S.D., Vysotskiy, S.V., 2003. Petrological diversity and origin of ments of K. Das become helpful in the final formulation of the man- ophiolites in Japan and Far East Russia with emphasis on depleted harzburgite. uscript. Critical comments of an anonymous reviewer and Dr. Hafiz Geol. Soc. Lond. Spec. Publ. 218, 597-617. Ur Rehman have also helped a lot to improve the scientific content Ishiwatari, A, Ozawa,K,Arai, S., Ishimaru, S.,Abe, N.,Takeuchi, M., 2016.Ophiolites and ultramafic rocks. In: Moreno, T., et al. (Eds.), The Geology of Japan. Geological Society, of this manuscript. This study was partly supported by Grant-in-Aids London, pp.223-250. for Scientific Research (no. 20540445 and 25400486: Y. Hayasaka) Isozaki, Y, Tamura, H., 1989. Late Carboniferous and early Permian radiolarians from the from the Ministry of Education, Culture, Sports, Science, and Tech- Nagato Tectonic Zone and their implication to geologic structure of the Inner Zone, Southwest Japan. Mem. Geol. Soc. Japan 33, 167-176 (in Japanese with English nology (MEXT) of Japan. abstract). Kagami, H, lizumi, S, Tainosho, Y., Owada, M, 1992 Spatial variations of Sr and Nd is References tope ratios of Cretaceous-Paleogene granitoid rocks, Southwest Japan Arc. Contrib. Mineral. Petrol. 112, 165-177. Aitchison,JC., Ireland, TR, Blake, M.C., Flood, PG., 1992. 530 Ma zircon age for ophiolit Kanazawa, T., Sager, W.W., Escutia, C., 2001. Explanatory notes. Proceedings of the Ocean from the New England orogen: oldest rocks known from eastern Australia. Geology Drilling Program, p. 46 Initial Reports 191. Katsube, A., Hayasaka, Y, Sakaguchi, A., Takahashi, Y, 2012. 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Kurokawa, K., 1985. Petrology of the Oeyama Ophiolitic Complex in the Inner Zone of chronological evidence for a Neoproterozoic ocean basin recorded in the Marlborough terrane of the northern New England Fold Belt. Aust. J. Earth Sci. 47, Southwest Japan. Series E (Geology & Mineralogy) No. 6Science Reports of Nigata University, pp. 37-113. 1053-1064. Dunkl, I, Mikes, T., Simon, K, von Eynatten, H, 2008.Brief introduction to the windows Ledneva, G.,Bazylev,BA, Lebedev, VV, Konnkova, NN., hwatari,A, 2012. Ub ir- program pepita: data visualization and reduction, outlier rejection, calculation of con age of gabbroids of the Ust'-Belaya Mafic-Ultramafic Massif (Chukotka) and its trace element ratios and concentrations from LA-ICP-MS data. In: Sylvester, P. (Ed.), interpretation. Geochem. Int. 50, 44-53. Laser Ablation ICP-MS in the Earth Sciences: Current Practices and Outstanding Is- Lindsley-Griffn, N., 1994. The Cambrian Trinity ophiolite and related rocks of the lower Paleozoic Trinity complex, northern California, U.S.A. In: Ishiwatari, A., et al. (Eds.), Sues. vol. 40. Mineralogical Association of Canada Short Course, pp. 334-340. Grimes, C.B., Joh, BE.,Klmen, PB, Mazdab, FK., Wooden, J.L, Cheadle, MJ, Hanghj Circum-Pacific Ophiolites. Proceedings of the 29th International Geological Congress K., Schwartz, JJ., 2007. Trace element chemistry of zircons from oceanic crust: a (Kyoto) Part D. VSP Publishers, The Netherlands, pp. 47-68. method for distinguishing detital zircon provenance. Geology 35, 643-646. Ludwig, K.R., 2003. User's Manual for Isoplot 3.00. A Geochronological Toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication No. 4, Berkeley, Harper, G.D., Saleeby, J.B., Heizler, M., 1994. Formation and emplacement of the Josephine ophiolite and the Nevadan orogeny in the Klamath Mountains, California-Oregon: U/ CA, p. 70. Pb zircon and 4Ar/39Ar geochronology.J. Geophys. Res. Solid Earth 99, 4293-4321. Lugmair, G.W., Marti, K, 1978. Lunar initial 143Nd/144Nd: differential evolution of the Hayasaka, Y., 1987.Study on the Late Paleozoic Early MesozoicTectonic Development f lunar crust and mantle.Earth Planet. Sci. Let.39, 349-357. Western half of the Inner Zone of Southwest Japan. vol. 27.Geological report of the Lus, W., MDougall ID, Hugh, L,2004. Age of the metamorphic sole of the Papuan U- Hiroshima University, pp. 119-204 (in Japanese with English abstract). tramafic Belt ophiolite, Papua New Guinea. Tectonophysics 392, 85101. Hayasaka, Y., Sugimoto, T, Kano, T., 1995. Ophiolitic complex and metamorphic rocks in Lyubetskaya, T., Korenaga,J., 2007. Chemical composition of Earth's primitive mantle and the Nimi-Katsuyama area, Okayama Prefecture. Excursion guidebook 102nd annual its variance: 1. Method and results. J. Geophys. Res. 112, B03211. meeting of the geological society of Japan. J. Geol. Soc. Jpn. 101 (Supplement), 71-87. Machi, S., Ishiwatari, A, 2010. Ultramafic rocks in the Kotaki area, Hida Marginal Belt, ceontral Japan: peridotites of the Oeyama ophiolite and their metamorphism. Hayasaka, Y., Moiseev, A.V., Sokolov, S.D., Ishiwatari, A., Machi, S., Ledneva, G.V., J. Geol. Soc. Jpn. 116, 293-308 (in Japanese with English abstract). Palandzhyan, S.A,Basylev,B.A., 2010. Methodology and philosophy for detrital zircon chronologyusing PMA,AICPMS nd SHRIMP,and outlineof results forthea Mankinen, EA.,Lindsly-Griffin, N, Griffn, J.R., 2002. Concordant paleolatitudes or zoic to Mesozoic complex in the Ust-Belaya Range,West Koryak Thrust and Fold Belt, Neoproterozoic ophiolitic rocks of the Trinity complex, Klamath Mountains, Califor- Far East Russia. Abstracts of the Russian-Japanese Workshop Symposium “Ophiolites nia. J. Geophys. Res. 107, 2254. and Related Complexes: Significance for Geodynamic Interpretations"June 15-16, Meert, J.G., Torsvik, T.H, 2003. The making and unmaking of a supercontinent: Rodinia Moscow, 2010, p. 11 2010. revisited Tectonophysics 375, 261-288. Herzig, C.T., Kimbrough, D.L., Tainosho, Y., Kagami, H, izumi, S., Hayasaka, Y., 1998. Late Nishimura, Y, Shibata, K., 1989.Modes of occurrence and K-Ar ages of metagabbroic rocks Cretaceous U/Pb zircon ages and Precambrian crustal inheritance in Ryoke granitoids, in the "Sangun metamorphic belt", Southwest Japan. Mem. Geol. Soc. Jpn. 33, Kinki and Yanai districts, Japan. Geochem. J. 32, 21-31. 343-357 (in Japanese with English abstract). Hofmann, AW., 2007. Sampling mantle heterogeneity through oceanic basalts: isotopes Nozaka, T, 2014. Metasomatic hydration of the Oeyama forearc peridotites: tectonic im- and trace elements. In: Holland, H.D., Turekian, K.K. (Eds.), Treatise on Geochemistry. plications. Lithos 184-187, 346-360. The Mantle and Core vol. 2. Pergamon, Oxford, pp. 1-44. 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Tsutsumi, Y., Yokyama, K, Miyawaki, R, Matsubara, S., Terada, K, Hidaka, H., 2010.Ags Tikhomirov, P.L., 2010. Age of plagiogranite from the Ust-Belaya Ophiolite complex of zircons in jadeitite and jadeite-bearing rocks of Japanese Islands. Bull. Nat. Mus. (Western Koryak Folded Zone): evidence from SHRIMP U-Pb dating of zircons. Nat. Sci. Ser. C 36, 19-30. Doklady Earth Sci. 434, 1362-1365. Wallin, E.T., 1990. Petrogenetic and tectonic significance of xenocrystic Precambrian zir- Toyohara, F, Yano, K., Doi, N., 1992. Thrust fault dividing the 500 Ma metagabbro from the con in Lower Cambrian tonalite, eastern Klamath Mountains, California. Geology 18, 170 Ma crystalline schists at the Cape Nomo, Nagasaki Peninsula. Kumamoto J. Sci. 1057-1060. Geol. 13, 49-59 (in Japanese with English abstract). Wallin, E.T., Metcalf, R.V., 1998. Supra-Subduction zone ophiolite formed in an exten- Tsujimori, T., 1999. 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1 Supplementary Information Groundwater oxygen isotope anomaly before the M6.6 Tottori earthquake in Southwest Japan Satoki Onda1, Yuji Sano1*, Naoto Takahata1, Takanori Kagoshima1, Toshihiro Miyajima1, Tomo Shibata2, Daniele L. Pinti3, Tefang Lan4, Nak Kyu Kim5, Minoru Kusakabe5 & Yoshiro Nishio6 1Atmosphere and Ocean Research Institute, The University of Tokyo, Kashiwa, Chiba 277-8564, Japan. 2Institute for Geothermal Sciences, Kyoto University, Beppu, Oita 874-0903, Japan. 3GEOTOP & Département des sciences de la Terre et de l’atmosphère, Université du Québec à Montréal, Montreal H3C 3P8, Canada. 4Department of Geosciences, National Taiwan University, Taipei, Taiwan. 5Korea Polar Research Institute, 26 Songdomirae-ro, Yeonsu-gu, Incheon 21990, Korea 6Graduate School of Integrated Arts and Sciences, Kochi University, Kochi 783-8502, Japan *Correspondence and requests for materials should be addressed to Y.S. (e-mail: ysano@aori.u-tokyo.ac.jp)
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Supplementary Figure 1. A relationship between δ18O and δD values of groundwater samples of ordinary period from Hakusan Meisui site. Fractionation of meteoritic water in the region (LMWL) is shown by a red dotted line with d=18 (ref. 22). MSS indicates precipitation data at Misasa. A solid line showes best fit by a least squre method with slope of 3.1. Error assigned to the simbol is 2σ.
3 Supplementary Figure 2. Temporal variations of anions in deep groundwater at the Hakusan Meisui site from September 2015 to July 2017. (a) Those of chloride (Cl-) and (b) sulfate (SO42-). Error assigned to the symbol is 2σ. Both contents show small inreases three month before the M6.6 earthquake.
4 Supplementary Figure 3. A relationship between δ18O and δD values of water samples recovered after frozen crushing experiments. Data of hot and cold blanks, 1500m and 240 m samples are shown with 2σ error. Evapolation line is calculated based on data of hot and cold blanks.
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Geographical distribution of3He/4He ratios in north Kyushu, Japan:
Geophysical implications for the occurrence of mantle-derived fluids
at deep crustal levels
Keika Horiguchi ⁎, Jun-ichi Matsuda
Department of Earth and Space Science, Graduate School of Science, Osaka University, Toyonaka, Osaka 560-0043, Japan
abstract article info
Article history:
Received 21 December 2011Received in revised form 15 December 2012Accepted 17 December 2012Available online 23 December 2012
Editor: L. Reisberg
Keywords:
HeliumIsotopeHot springFluid
Beppu-Shimabara Graben
KyushuWe measured3He/4He ratios in hot spring gases collected from sites in north Kyushu, Japan that represent an
island arc setting associated with the subduction of the Philippine Sea plate beneath the Eurasian plate. Our
results revealed relatively high3He/4He ratios (about 4 to 8 Ra) in the Beppu-Shimabara Graben (a continu-
ation of the Okinawa trough) and relatively low ratios (about 0.1 to 2 Ra) in areas located to the north and
south of the graben. Slightly elevated ratios (about 3 to 4 Ra) were observed in areas further north of the gra-
ben. North Kyushu hosts unusual He isotopic signatures considering that its backarc location would predicthigher
3He/4He ratios than the low ratios observed here. Instead of re flecting a regional tectonic signature,
the geographic distribution of3He/4He ratios appears to vary with seismic velocity perturbations as detected
at 20 –30 km depth. The correspondence between these two parameters suggests that3He/4He ratios in this
case relate to the spatial distribution of fluids at the base of the crust.
© 2012 Elsevier B.V. All rights reserved.
1. Introduction
Helium is a well-known tracer for volatiles originating from
the mantle. In a typical island arc setting like northeastern Japan,
3He/4He ratios are relatively low within the forearc and higher within
the volcanic arc (volcanic front and backarc; Sano and Wakita, 1985 ).
Elevated3He/4He ratios observed within the volcanic arc relate to
rising mantle material (itself having high3He/4He ratios), whereas
the low3He/4He ratios within the forearc relate to the addition of
radiogenic4He from U and Th decay occurring in continental crustal
material. Kyushu Island is one of the four major Japanese Islands
and lies just southwest of the main island ( Fig. 1 ). Kyushu developed
through an unusual set of tectonic events as described in the next sec-tion. Helium isotopic ratios in Kyushu have been reported primarily
from volcanic gases (e.g., Nagao et al., 1981; Sano and Wakita, 1985;
Kita et al., 1993 ). More recent studies have measured the helium iso-
tope ratios from hot spring waters and groundwaters in the region(Umeda et al., 2007; Mahara and Kitaoka, 2009 ). These studies sam-
pled relatively few hot spring sites relative to studies that addressedHe signatures in other regions of SW Japan ( Matsumoto et al., 2003;Sano and Nakajima, 2008; Sano et al., 2009 ). Previous studies have
thus failed to present consistent patterns in helium isotopic signatureswithin the Kyushu district. A study of
3He/4He ratios that directly
addresses the peculiar tectonic setting of Kyushu can resolve someof these uncertainties.
In this study, we measured
3He/4He ratios of water samples from
hot springs located in the south Fukuoka and north Kumamoto pre-
fectures. This area covers the Beppu-Shimabara Graben and regions
immediately to the north and south ( Fig. 2 ). The research seeks to
establish the surface spatial distribution of helium isotopic signaturesand interpret patterns according to tectonic features in the study area.
This work is also the first report concerning
3He/4He ratios of an active
continental rift valley (extensional feature) of Japan. We believe sucha geochemical approach should be useful for understanding the mech-
anisms of big earthquakes when combined with the geophysical
approach such as seismology and geomagnetism.
2. Tectonic setting
The Japanese Islands consist of two distinct blocks: northeastern
(NE) and southwestern (SW) Japan. The Itoigawa –Shizuoka tectonic
line divides NE Japan from SW Japan. NE Japan is a typical island arc
located at the boundary between the Paci fic plate and the North
American/Okhotsk plate, which converge at a rate of about 8 –9c m / y r
(Fig. 1 ; e.g., DeMets et al., 1990, 1994; Seno et al., 1996 ). LargeChemical Geology 340 (2013) 13 –20
⁎Corresponding author at: Crustal Fluid Research Group, The Institute of Geology and
Geoinformation, Geological Survey of Japan, AIST, 1-1-1 Higashi, Tsukuba, Ibaraki,
305-8567, Japan.
E-mail address: keika@ess.sci.osaka-u.ac.jp (K. Horiguchi).
0009-2541/$ –see front matter © 2012 Elsevier B.V. All rights reserved.
http://dx.doi.org/10.1016/j.chemgeo.2012.12.008
Contents lists available at SciVerse ScienceDirect
Chemical Geology
journal homepage: www.elsevier.com/locate/chemgeo
earthquakes repeatedly occur at the boundary between the continental
and oceanic plates while inland earthquakes occur at shallow crustal
levels in association with local compressional/tensional deformation.
SW Japan is located at the boundary between the Philippine Sea plate,
a relatively young and warm body, and the Eurasian plate, both of
which converge at a rate of ~4.0 –6.5 cm/yr ( Shiono and Sugi, 1985;
Seno et al., 1993; Miyazaki and Heki, 2001 ). Large interplate earth-
quakes occur in SW Japan with similar magnitude to those in NE
Japan, and with a recurrence interval of about 100 –200 years ( Ando,
1975 ).
Kyushu Island is located at the intersection of the SW Japan arc
and the Ryukyu arc. The Philippine Sea plate subducts beneath the
Eurasian plate along the Ryukyu trench and the Nankai Trough
(Fig. 1 ). The subduction angle at the Ryukyu trench exceeds that
of the Nankai Trough. The former subduction angle is close to orthog-
onal, which is seismically detected to depths of up to 150 –180 km
(Nakajima and Hasegawa, 2007 ). The Okinawa Trough extends to
the west of the Ryukyu trench as a backarc basin of the Ryukyu arc.The tectonic setting of Kyushu Island is unusual for several reasons
(Fig. 2 ). Its location and tectonic setting suggest that it relates to
the subduction of the Philippine Sea plate, but triangulation and
trilateration surveys have demonstrated that extensional tectonics
operate in the area. The crust in central Kyushu is extending in a
N–S direction with spreading centered at the Beppu-Shimabara Gra-
ben ( Tada, 1984, 1985, 1993 ). Gravity anomalies show crustal
depression and uplift of the Moho discontinuity (crustal thinning)
beneath the Beppu-Shimabara Graben. The spreading subsidence
rates are estimated to be 1.4 cm/yr and 2.5 cm/yr, respectively ( Tada,
1993 ). The Beppu-Shimabara Graben is interpreted as the northeast-
ern extension of the Okinawa Trough, itself interpreted as a typical
backarc basin. The inland location of this backarc feature is unusual.
Central Kyushu experiences frequent shallow earthquakes as well as
volcanism resulting from dehydration processes along the descending
Philippine Sea plate associated with melting of the mantle wedge.
Upwelling of molten materials from backarc opening has also con-
tributed to the Unzen volcanism in the Beppu-Shimabara Graben130° 135° 140° 145°25°30°35°40°45°
Philippine Sea PlatePacific PlateEurasian Plate
Volcanic front
Nankai Trough
Kyusyu-Palau ridge
Volcanic frontOkinawa TroughNorth American Plate
(Okhotsuk plate)
Itoigawa-SizuokatectoniclineJapan Trench
~4.0-6.5 cm/y~8-9 cm/y
Ryukyu TrenchKyushu island
Fig. 1. Tectonic setting of the study area and the surrounding Eurasian, Philippine Sea, Paci fic and North American (Okhotsuk) plates. The study area lies to the north of the Okinawa
Trough, to the west of the Nankai Trough and Ryukyu Trench, and southwest of the Japan Trench and Itoigawa –Shizuoka tectonic line, etc.14 K. Horiguchi, J. Matsuda / Chemical Geology 340 (2013) 13 –20
(Wang and Zhao, 2006 ). North Kyushu thus hosts two different types of
volcanism.
3. Sampling and noble gas analyses
We collected 14 water samples from hot springs in the study area
(Fig. 2 ; solid circles). The sites sampled for He analyses by previous
studies are shown as open circles ( Sano and Wakita, 1985; Umeda
et al., 2007 ). Our samples were collected using 100 cm3pyrex glass
containers with vacuum valves at both ends (e.g., Matsumoto et al.,
2003; Horiguchi et al., 2010 ). In many cases, water from the hotsprings was pumped to the surface, allowing us to collect samples
directly from an output tap. In other cases hot spring water was col-
lected from a storage tank which we sampled using a clean syringe.
We avoided exposure to ambient air as much as possible during the
sampling process. For the “Ajisainoyu ”sample (#10 in Fig. 2 ), we col-
lected water from a hot spring bathing facility because the spring
source was geographically inaccessible. We performed temperature
measurements on site or collected temperature data from literature
sources. At the “Tamana ”locality (#11 and #12 in Fig. 2 ), samples
were taken from two wells (Nos. 1 and 3) extending to different
depths in the subsurface in order to compare results.
↑↑
↑
↑↑
↑
129° 130° 131° 132°31°32°33°34°
↑
↑
↑ ↑↑
↑
UnzenVolcanic front
Kuju
AsoTsurumi/Garan
Yufu
Kirishima
SakurajimaFukuoka
Saga
KumamotoOita
MiyazakiNagasaki
KagoshimaNagasaki peninsulaChikushi Mts.
ShiroishiSaga plain
Miyazaki plain1 2
6
7
8
9104
5 11
1213314Beppu-Shimabara
Graben
Beppu-Shimabara Graben
Fig. 2. Tectonic setting of north Kyushu. The double lines show rift zones (Beppu-Shimabara Graben) and the single lines show associated shear faults. Activ e volcanoes are shown
as black triangles and the volcanic front is shown as a cross pattern between dashed lines. Thin lines indicate active faults. Solid circles represent sampling locations used in the
present study and open circles indicate those of previous studies (e.g., Sano and Wakita, 1985; Umeda et al., 2007 ). Numerical values next to sampling localities correspond to
those given in Table 1 .15 K. Horiguchi, J. Matsuda / Chemical Geology 340 (2013) 13 –20
Gases dissolved in water samples were extracted by releasing the
water into an evacuated glass container connected to a preparation
and puri fication line for noble gas analyses. The3He/4He ratios were
measured along with4He and20Ne concentrations using a VG5400
Noble Gas mass spectrometer at Osaka University. The helium isoto-
pic ratios measured in samples were calibrated with the arti ficial
internal He Standard of Japan (HESJ; Matsuda et al., 2002 ). Further
details concerning the noble gas measurement techniques used
here can be found in Horiguchi and Matsuda (2008) andMatsumoto
et al. (2003) . The analyses did not require blank corrections due to
negligible levels of the measured noble gases observed in procedural
blanks relative to the measured peaks for unknown samples.
4. Results and discussion
Table 1 lists results of the noble gas analyses. Fig. 3 shows concen-
trations of4He and20Ne relative to hot spring temperatures. Concen-
trations of4He varied more than those of20Ne with the latter giving
consistently lower values except in the case of the “Ajisainoyu ”
(#10) sample. Collection from a surface flow is likely to have caused
atmospheric contamination in the case of the “Ajisainoyu ”sample.
Data from the two “Tamana ”wells (#11 and #12) show similar4Heand20Ne concentrations ( Fig. 3 ) demonstrating that the different
collection depths did not affect gas ratios.
Fig. 3 also shows reference4He and20Ne concentrations dissolved
in air-saturated water (ASW) at 1 atm ( Weiss, 1971 ). The4He con-
centrations measured in the hot spring samples exceed those in
ASW by about 1 to 3 orders of magnitude, whereas the20Ne con-
centrations are only about 5 times higher than those observed in
ASW with the exception of samples 1 and 14 which display sub-
atmospheric20Ne concentrations. Assuming that noble gases behave
in an ideal fashion, solubility is proportional to the subsurface pres-sure. The Ne observations therefore indicate that gases dissolve in
the water at pressures of about 5 atm in the subsurface. Unlike the
20Ne data, the elevated concentrations of4He appear to re flect a
different mechanism.
Fig. 4 shows the measured3He/4He ratio ((3He/4He) m) plotted
against the20Ne/4He ratio. We plotted (3He/4He) mversus20Ne/4He
instead of the more typical (3He/4He) mversus4He/20Ne construct
used in many papers (e.g. Sano and Wakita, 1985 ). The use of the
common4He denominator presents mixing relationships between
distinct components as straight lines ( Matsuda and Marty, 1995 ).
The3He/4He ratio of air (1.4×10−6), is de fined as 1 Ra while theTable 1
Sampling site, temperature,4He/20Ne and3He/4He ratios in the Kyushu district in SW Japan.
No. Sample name Date Location (N, E) Temperature
(C°)4He
(cm³STP/g)20Ne
(cm³STP/g)(3He/4He) m
measured (Ra)4He/20Ne (3He/4He) c
corrected (Ra)(N ˚)( E ˚)
1 Katanose 2010.8.9 33.34 130.69 27.5 7.31E −06 6.05E −08 2.939±0.035 120.7 2.944±0.035
2 Harazuru 2010.8.9 33.35 130.78 45.6 1.74E −05 7.37E −07 3.686±0.045 23.62 3.723±0.046
3 Kuroki 2010.8.9 33.18 130.68 33 9.62E −07 5.26E −07 0.266±0.012 1.828 0.112±0.015
4 Hirayama 2010.8.9 33.05 130.66 23 2.80E −05 5.43E −07 0.647±0.010 51.46 0.645±0.010
5 Kikuchi 2010.8.10 32.99 130.81 47 1.37E −05 7.43E −07 1.359±0.020 18.476 1.366±0.020
6 Otsu 2010.8.10 32.87 130.95 58.1 2.22E −05 5.12E −07 4.795±0.044 43.479 4.823±0.044
7 Suizenji 2010.8.10 32.79 130.74 38.7 3.74E −05 6.78E −07 3.789±0.042 55.221 3.805±0.042
8 Mifune Kannon 2010.8.10 32.72 130.80 40.4 2.66E −05 7.16E −07 1.950±0.028 37.241 1.959±0.028
9 Hama no Yu 2010.8.10 32.68 130.99 31.4 3.16E −06 5.77E −07 0.144±0.005 5.488 0.091±0.005
10 Ajisainoyu 2010.8.10 32.69 130.60 26.5 2.32E −07 3.98E −07 1.192±0.058 0.582 1.424±0.129
11 Tamana (Well 1) 2010.8.11 32.94 130.56 41.8 2.18E −06 5.91E −07 0.796±0.012 3.687 0.776±0.014
12 Tamana (Well 3) 2010.8.11 32.94 130.56 44 2.80E −06 7.69E −07 0.742±0.017 3.644 0.717±0.019
13 Ueki 2010.8.11 32.96 130.72 51 1.68E −05 5.91E −07 1.766±0.026 28.435 1.775±0.026
14 Funagoya 2010.8.11 33.18 130.51 20 3.12E −07 3.65E −08 2.979±0.089 8.56 3.055±0.092
10-810-710-610-5
10 20 30 40 50 604He
20Ne
Temperature (°C)4He content in ASW20Ne content in ASW144
1
9
1037
8
111225136
144
10
1937811
1225136Concentration (cm3 STP/g)10-4
Fig. 3.4He (open circles) and20Ne (closed circles) concentrations plotted against the
temperatures from hot spring water samples. Numerical values are sample localitiesindicated in Fig. 2 andTable 1 . Dotted lines show
4He and20Ne reference concentra-
tions in air-saturated water ( Weiss, 1971 ).0.001.002.003.004.005.00
0 0.5 1 1.5 2 2.5 3 3.5
20Ne/4HeAir106
7
2
14
1
8
13
5
11
12
9 34to MORB source (8 Ra)
to Crustal source
(0.001 Ra)
(3He/4He) m (Ra)
Fig. 4. A diagram of (3He/4He) magainst20Ne/4He. Values for (3He/4He) mare given in
Ra (1 Ra=1.4×10−6;R ar e f e r st ot h e3He/4He ratio relative to the atmosphere value).
The closed circle indicates the atmospheric value (3He/4He=1 Ra and20Ne/4He=3.14).
The diagram also shows MORB (8 Ra) and radiogenic (0.001 Ra) end-member composi-tions. Both end-members have a very low
20Ne/4He ratio (roughly 1/5000). Numerical
values give sample localities indicated in Fig. 2 andTable 1 .16 K. Horiguchi, J. Matsuda / Chemical Geology 340 (2013) 13 –20
20Ne/4He ratio of air is 3.14 (4He/20Ne=0.318). The20Ne/4He ratio of
mantle and crustal sources is very low (~1/5000). The3He/4He ratio
of mantle sources is about 8 Ra whereas crustal sources have3He/
4He ratios of about 0.001 to 0.02 Ra. The measured data fall within
the ternary space de fined by mantle, crustal and atmospheric sources
(Fig. 4 ). The data are consistent with other He isotope data from hot
springs in Japan.The atmospheric He component detected in samples may occur
in-situ or may enter in samples upon exposure to the atmosphere
during sampling. We performed the following calculation to subtract
the atmospheric He component:
ð3He=4HeÞc¼ð1=5000−1=0:318Þ=ð1=ð4He=20NeÞm−1=0:318Þ
/C2ð ð3He=4HeÞm−1Þþ1↑↑↑↑
↑↑↑
↑
↑ ↑↑
↑Fukuoka
Saga
KumamotoOita
MiyazakiNagasaki
KagoshimaNagasaki peninsulaChikushi Mts.
ShiroishiSaga plain
Miyazaki plain
5.68
129° 130° 131° 132°31°32°33°34°
2.94 3.72
1.37
4.82
3.81
1.960.091.786.52
5.85
7.074.56
0.47
0.68
0.79
0.29
0.780.3
0.21
0.09
0.31
0123456
3He/4He [R atm]4.335.250.72
0.780.653.06 0.116.79
6.43
1.42
6.04
6.17
3.795.24
2.51Volcanic front
Fig. 5. Geographical distribution of the corrected3He/4He ratios ((3He/4He) c) relative to the tectonic setting of Kyushu Island. The3He/4He ratios come from this study (large
circles) and from previous studies ( Sano and Wakita, 1985; Umeda et al., 2007 ; smaller circles). Circles are color coded according to3He/4He ratios in Ra, with numerical Ra values
adjacent.17 K. Horiguchi, J. Matsuda / Chemical Geology 340 (2013) 13 –20
where (3He/4He) cis the corrected3He/4He ratio and the subscript
“m”indicates measured values. The (3He/4He) cis extrapolated from
the y-axis of the mixing line connecting the data point and atmo-
spheric values in Fig. 4 . Strictly speaking, (3He/4He) ccorresponds to
20Ne/4He=1/5000, but since this value is so close to zero, we can
treat it as the y-intercept. The calculated (3He/4He) cvalues re flect
only the mixture of mantle and crustal helium, as given in Table 1 .
5. Geographical distribution of3He/4He ratios in north Kyushu
The (3He/4He) cdata obtained in this study are shown in map view
along with data from previous studies ( Fig. 5 ). High3He/4He ratios
in the Beppu-Shimabara Graben have also been reported by Kita
et al. (1993) . These workers found that the N 2/Ar ratios in Kyushuwere an order of magnitude lower than those of NE Japan due to a
lesser contribution from subducted sediment in Kyushu magma
sources. This interpretation has not been tested using helium isotopic
signatures.
The low3He/4He ratios observed in north Kyushu are unusual
considering the area's backarc location relative to the volcanic front.
3He/4He ratios are typically low in the forearc region and higher in
and around the volcanic front and backarc region (e.g., Sano and
Wakita, 1985 ). The3He/4He ratios are high (about 4 to 8 Ra) within
the Beppue-Shimabarta Graben but lower (about 0.1 to 2 Ra) in
areas to the north and south in spite of their general location within
the backarc ( Fig. 5 ). Slightly elevated3He/4He ratios (about 3 to 4
Ra) were also detected in areas of Fukuoka prefecture, north of the
graben.
Velocity perturbation (%)-10 -5 0 10 5129° 130° 131° 132°31°32°33°34°
0123456
3He/4He [Ratm]20kmVp
Fig. 6. Comparison of the observed3He/4He ratios and velocity perturbation for V P(P-wave velocity; Saiga et al., 2010 ). White circles indicate epicenters of events used in seismic
imaging ( Saiga et al., 2010 ).18 K. Horiguchi, J. Matsuda / Chemical Geology 340 (2013) 13 –20
Mahara and Kitaoka (2009) also reported several low3He/4He
ratios ( b1 Ra) from central Shiroishi on the Saga plain (a few tens
of kilometers west of our sampling area). These observations indicate
that the groundwater in central Shiroishi may have accumulated
radiogenic He through mixing with fossil pore water from imperme-
able marine clay layers. Shiroishi experiences severe land subsidence
due to over-pumping of groundwater for rice cultivation, which mayexplain the entrainment of deeply sourced fossil water isolated from
the regional groundwater circulation ( Mahara and Kitaoka, 2009 ).
Our sampling area lies on the eastern side of Shiroishi in an area
where land subsidence has not been reported ( Fig. 5 ). The low
3He/4He ratio observed in our study is therefore not apparently due
to radiogenic4He accumulated in fossil water.
Low3He/4He ratios (less than 1 Ra) are unusual within a backarc
setting but also occur in NE Japan ( Horiguchi et al., 2010 ). For exam-
ple,3He/4He ratios are less than 1 Ra in and around the border area
between the Iwate and Akita prefecture (39 –39.5° N). The geographic
distribution of helium isotope ratios in these areas coincides with
seismic velocity and attenuation of seismic wave distributions in the
uppermost mantle, and re flects the distribution of melts in the upper-
most mantle ( Horiguchi et al., 2010 ).
Recently, Xia et al. (2008) and Saiga et al. (2010) published a
high-resolution three-dimensional image of velocity structure in
Kyushu. Saiga et al. (2010) compared the velocity anomaly distribu-
tion with the crustal gravity anomalies and interpreted them as evi-
dence of low-density material at the base of the crust. Low-velocity
anomalies exist beneath the Beppu-Shimabara Graben and northeast-
ern Kagoshima. High-velocity anomalies meanwhile exist beneath
the Miyazaki plain, the Nagasaki peninsula and Chikushi mountains
(Figs. 5 and 6 ). The low3He/4He ratios observed in our study corre-
spond to a region of high seismic velocity at about 20 –30 km depth
beneath areas north and south of the Beppu-Shimabara Graben,
as described in both Xia et al. (2008) and Saiga et al. (2010) . The
high3He/4He ratios (about 4 to 8 Ra) measured within the Beppu-
Shimabara Graben correspond to a low velocity anomaly area.
Fig. 6 compares3He/4He ratios with velocity perturbation for V P
(P-wave velocity) at 20 km depth as reported in Saiga et al. (2010) .
This comparison shows that the geographic distribution of high
and low3He/4He ratios agrees well with that of the respective low
and high seismic velocity regions. Slightly elevated3He/4He ratios
of about 3 to 4 Ra were observed in the north corner of the study
area, which overlie a low-velocity region. The origins of the low-
velocity zones beneath the Beppu-Shimabara Graben and its surround-
ings are not well understood, but it is likely that the low-velocity
zone at the base of the crust in north Kyushu is due to upward migra-
tion of mantle-derived fluids which would also have high3He/4He
ratios. The water supplied by dehydration of the subducting slab
reaches the upper crust via entrainment and upwelling through the
mantle wedge, a body characterized by lower relative seismic veloc-
ity and high attenuation in NE Japan ( Hasegawa et al., 2005 ). Regions
that do not host upwelling fluid-rich zones exhibit low3He/4He ra-
tios. We interpret these lower values to simply re flect regional crustal
values.
These observations indicate that helium isotopes can record the
presence of deep fluid-rich zones associated with regional tectonic
configurations. We conclude that the low3He/4He ratios in areas to
the north and south of the Beppu-Shimabara Graben re flect the
absence of deep mantle-derived fluids at the base of the crust. Mantle
fluids are detected within the Beppu-Shimabara Graben but are not
widely distributed throughout north Kyushu in spite of its backarc
location.
6. Conclusions
High3He/4He ratios (about 4 to 8 Ra) were observed in the
Beppu-Shimabara Graben, whereas low3He/4He ratios (about 0.1 to2 Ra) were observed in areas to the north and south of the graben.
Slightly elevated ratios (about 3 to 4 Ra) are also observed in areas fur-
ther north of the graben. The spatial distribution of these3He/4He ratios
is inconsistent with the region's backarc setting, a tectonic environment
that would predict3He/4He ratios greater than the atmospheric value,
similar to those observed in areas of NE Japan. We interpret the high
and low3He/4He ratios as re flecting respective low- and high-velocity
regions occurring at about 20 –30 km depth beneath the sampling re-
gion. This interpretation suggests that high3He/4He ratios relate to
the presence of deep mantle-derived fluids at the base of the crust.
Acknowledgments
We thank Dr. Laurie Reisberg (handling editor), Dr. Daniele Pinti
and one anonymous reviewer for constructive comments and
thoughtful suggestions that greatly improved the manuscript. We
also thank individual owners of hot springs for their cooperation dur-
ing sampling campaigns and Mr. Philip Nguyen for the help with
experimental procedures. This work was supported by a JSPS (Japan
Society for the Promotion of Science) Research Fellowship for Young
Scientists (No. 21-1377). Some of the figures in this paper were drafted
using GMT (Generic Mapping Tools) software written by Wessel and
Smith (1991) .
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Journal of Volcanology and Geothermal Research 433 (2023) 107739 Contents lists available at ScienceDirect 二 VOLCANOLOGY Journal of Volcanology and Geothermal Research ELSEVIER journal homepage: www.journals.elsevier.com/journal-of-volcanology-and-geothermal-research Groundwaters and deep-seated fluid circulation around Aso Volcano, Southwest Japan, revealed by multivariate statistical analysis of the geochemical data Hikaru Iwamori a *, Hitomi Nakamura b, Noritoshi Morikawa b, Masaaki Takahashi b, Akihiko Inamura , Satoru Haraguchi ?, Tatsuji Nishizawa , Shuhei Sakata a aEarthquake Research Institute, The University of Tokyo, Tokyo,Japan Geological Survey of Japan,National Institute of Advanced Industrial Science and Technology(AIST),Tsukuba,Japan MountFjhstinahitlventujhid ARTICLEINFO ABSTRACT Keywords: To understand the fluid circulation under volcanoes and their surrounding area, geochemical compositions of Groundwater pe no sunu dr ms u nsna nn a w sds snnm naus pe rmi o Deep-seated fluid mineral spring waters around Aso and Kuju Volcanoes, have been analyzed for 12 major solutes as well as Cluster analysis oxygen-hydrogen and helium isotopic ratios. To extract independent processes and sources involved in formation Geographical distribution of the groundwaters, whitened data-based k-means cluster analysis (KCA) has been applied to the data set of 12 Aso Volcano Kuju Volcano major solutes. Eight clusters have been identified by KCA, which were mapped geographically and were aa pn nu anm d p an s i m pn geographical mapping, a concentric zonal structure of the clusters emerges in the Aso caldera, which resulted broadly from two fluid cycling systems from caldera rim and central cones. In the entire study area including Aso and Kuju Volcanoes, the mapping also shows geographical provenance, and the distribution of each cluster shows characteristics related to geological features, such as upwelling of deep-seated fluids along the prominent tec- tonic lines. On the other hand, the Piper diagram-based classification is significantly different from the KCA results and exhibits neither a concentric zonal structure within the Aso caldera nor a regional geographical provenance. These results suggest that multivariate statistical analysis such as KCA is potentially useful for investigating the origins of groundwaters and shows the advantages over classical Piper diagram-based evaluations. 1. Introduction et al., 2014). Thus, combination of spatial distribution and chemical characteristics of groundwater will provide useful information con- Distribution and chemistry of groundwater, including hot and min- cerning flow, processes, and sources of groundwater beneath the vol- eral spring waters, within and around volcanoes are important in terms canic area as well as its surrounding area. of several aspects. For instance, the extent and locations of hydrothermal This study presents geochemical data of groundwater from the cen- discharges may control phreatic/phreatomagmatic eruptions (Stix and tral Kyushu, Southwest Japan, including Aso Volcano and Kuju Volcano. Moor, 2018) and a catastrophic volcanic collapse (Lopez and Williams, The geochemical data are statistically analyzed by an unsupervised 1993; Hurwitz et al., 2007). The chemistry of groundwater may work as machine learning type-approach (e.g., cluster and independent compo- a useful proxy to identify geochemical components derived from nent analyses) to discuss independent processes and sources that have magmas and fluid processes within the volcanic edifice (e.g., Morikawa controlled the groundwater chemistry. Aso Volcano is an active volcano et al., 2008a; Yamada et al., 2011; Ohwada et al., 2012) as well as deep- that has a large caldera (18 × 25 km in diameter) located in the middle seated fluid components circulating in a subduction zone scale, e.g.. part of Kyushu Island of the southwestern Japan arc (Fig. 1). The main those derived from a subducting slab (Kusuda et al., 2014; Nakamura activity of Aso Volcano started about 0.8 Ma and has continued to the * Corresponding author. E-mail address: hiwamori@eri.u-tokyo.ac.jp (H. Iwamori). https://doi.org/10.1016/j.jvolge0res.2022.107739 Received 21 June 2022; Received in revised form 10 December 2022; Accepted 17 December 2022 Available online 20 December 2022 0377-0273/? 2022 Elsevier B.V.All rights reserved. H. Iwamori et al Journal of Volcanology and Geothermal Research 433 (2023) 107739 present, including the four caldera-forming eruptions between 0.27 and Oxygen isotope ratio (818o) were measured with isotope-ratio mass 0.09 Ma (Furukawa et al., 2009). Beneath this area, the Philippine Sea spectrometers (Delta Plus, Thermo Fisher Scientific Inc.). CO2 generated Plate is subducting northwestward from the westernmost part of the by the automated H20-CO2 equilibration method was applied for 8180 Nankai trough with a convergence velocity of ~5 cm/year (Seno et al. analysis. Analytical errors were ± 0.1%o. Hydrogen isotope ratio (SD) were measured with isotope-ratio mass spectrometers (Delta V Advan- to supply aqueous fluids to the overlying mantle wedge and arc crust to tage, Thermo Fisher Scientific Inc.) or analyzed with a cavity ring-down induce arc magmatism at Aso and Kuju Volcanoes (Zhao et al., 2000; spectrometer (L2120-i, Picarro Inc.) with analytical errors of ±0.6%. Iwamori, 2007). The H2 generated by the H2O-H2 reduction method by 800 °C metal In this study, we analyzed the compositions of groundwaters chromium (H-device, Thermo Fisher Scientific) was applied for mass sampled over the area of about 120 km east-west and 80 km north-south spectrometric analyses and analytical errors were ± 1%o. surrounding the Aso caldera in central Kyushu (Fig. 1). In particular, we aim to distinguish the volcanic and deep-seated fluid components from tubes, and free gas was sampled in glass bottles if bubbling gas was those related to meteoric water circulation, and ultimately to under- visually observed. Helium and Ne concentrations and helium-isotopic stand (1) fluid processes associated with magmatic activity under the ratios were measured using a static noble gas mass spectrometer Aso caldera, and (2) the large-scale fluid circulation from the subducting (Micromass MM5400). Technical details are described by Morikawa plate to the surface, including Aso and Kuju Volcanoes and the sur- et al. (2008b). Ten repeated analyses of *He/*He and noble gas con- rounding non-volcanic area, in central Kyushu (Fig. 1). To distinguish centrations for air-saturated water was within 3% average value (1o). the different sources and processes that produced the groundwaters and their chemistry, multivariate statistical analyses will be applied to the these groundwaters, we used a multivariate statistical analysis for this compositional data for major dissolved elements, particularly the cluster high-dimensional data set. The statistical methods used is “whitened analysis. Together with considerations based on the isotopic ratios of data-based k-means cluster analysis" (hereafter referred to as KCA) and oxygen, hydrogen and helium, the origins of the groundwaters were data (Iwamori et al., 2017). Of the standard multivariate analyses, from relatively shallowly circulating waters. Finally, the statistically principal component analysis (PCA) is common to find the principal identified clusters are compared with the Piper diagram-based classifi- components (PCs) that accounts for the sample variance most effec- cation, showing the utility of the statistical approach. tively. The extracted PCs are uncorrelated with each other (in other words, orthogonal) but are not necessarily independent. When the data 2. Data and Methods constitute a (joint) non-Gaussian distribution, PCA always fails to extract the independent features in the data (Hyvarinen et al., 2001). We examined the high-dimensional composition data set established Instead, KCA is suitable for data showing non-Gaussian distribution. As by Takahashi et al. (2022), which consists of 590 water samples with 12 will be shown in the next chapter, the concentration data used in this major solute concentrations of Na, K, Mg, Ca, Li, Cl, SO4, HCO3, F, NO3, study show non-Gaussian distributions and KCA is used in the following Br, and total C (Supplementary material, Table S1). analysis. Here, statistical analysis was performed using the same pro- The analytical procedures are as follows. Concentrations of Na, K, cedures and methods as in Iwamori et al. (2020), involving the pre- Mg, Ca, Li, Cl, SO4, F, NO3 and Br, were measured by an ion chroma- processing the data by logarithmic transformation and whitening (i.e., tography system (Thermo Fisher Scientific DXi-500 and ICS-2100). normalizing the data along the eigenvectors by the square root of the e o dn rs n sm eigenvalues to make the new transformed variables uncorrelated with 4.8 using an automatic potentiometric titrator (Kyoto Electronics equalized variance of unity). This method has been shown to be suitable Manufacturing AT-510). The amount of total carbon (abbreviated as for extracting independent features in the groundwater solute concen- "total C" or *TotC"), HCO3 and CO3- concentrations were determined by tration data, including the Arima-type brine and the meteoric water the thermodynamic calculation using the alkalinity, temperature and cycling (Iwamori et al., 2020). pH. 33.2 SHIKOKU 33 32.8 32.6 PHILIPPINE SEA 20km PLATE 130.4 130.6 130.8 131 131.2 131.4 131.6 Fig. 1. Sampling points of groundwater and spring/river water (left, red points) in the study area including Aso and Kuju Volcanoes, and the index map of the study area (right) with the location and size of the Aso caldera, central Kyushu, beneath which the Philippine Sea Plate subducts with the velocity of ~5 cm/year (white arrow) from the Nankai Trough. The gray scale coding (left) indicates the altitude in meter, based on SRTM15_plus ver.2.1. In the left figure, the thin solid line represents the coastline (O m above sea level) and the broken curved line outlines the Aso caldera. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article). H. Iwamori et al. Journal of Volcanology and Geothermal Research 433(2023) 107739 3. Results of KCA the side of lower K concentration at a constant Na concentration). These trend and relation are readily seen in Fig. 3(b), as an up-and-down 3.1.Characteristicsofmajorsolutesofclusters pattern in the K concentration diagram. Different patterns are seen depending on the solutes; e.g., Mg and Ca show broadly positive trends Eigenvalue analysis of compositional data (12 analytes) for a total of (Figs. 2(b) and 2(c), and Figs. 3(c) and 3(d)), but Cluster 6 (purple 590 samples revealed that 84.3% of the sample variance is explained by square) is off significantly below the trends and is also characterized by the three eigenvectors with the three largest eigenvalues (60.4%, 13.4%, the highest mean F and the lowest mean NOg concentrations (Figs. 3(i) 10.4%). The fourth and the subsequent eigenvectors account for <5% of and 3(j)). In the SO4 diagrams (Figs. 2(f) and 3(g)), again Cluster 6 is off below the main trend and Cluster 8 plots even below. Cluster 8 shows spanned by the three major eigenvectors, eight quadrants emerge in the the lowest mean for SO4 (Fig. 3(g)) although it is characterized by the whitened space, which gives a criterion for determining the number of highest mean concentrations in many other solutes (Na, Li, Cl, HCO3, Br, clusters (Iwamori et al., 2017). Based on these, the samples were clas- and Total C) as the densest cluster. On the most diluted side, Clusters 1 sified into eight clusters in this study. and 2 (black and red) always constitute confined positive trends for all Fig. 2 shows the variation diagrams between Na and other solute the 12 solutes (Figs. 2 and 3). Cluster 3 overlaps these trends of Clusters concentrations, with different symbols and colors for the individual 1 and 2, except for K). Cluster 4 constitutes an extension of the trends by clusters from No. 1 to 8. The clusters are numbered according to the Clusters 1 and 2 in most solutes with some deviation to lower concen- mean Na concentration of each cluster as in the top-left diagram of tration sides in HCO3, NO3, and Total C. The deviation of Cluster 4 in Fig. 3. In the binary diagrams of Fig. 2, the individual clusters form the NO3 is particularly remarkable to form an overall broad trend with clouds of data points. The clouds overlap with each other in these binary Clusters 5 to 8 with a negative slope. Cluster 7 is characterized by the diagrams due to projection of the 12-dimensional data on to two- highest mean K, Mg, Ca, and SO4, and Cluster 5 broadly shows inter- dimensional planes, yet Figs. 2 and 3 are useful to show the composi- mediate ranges between Clusters 4 and 7 in all the solutes except HCO3 tional characteristics of individual clusters. (Fig. 3(h)). For instance, the Na—K variation (Fig. 2(a)) indicates that Clusters 1, The distribution of data points in Fig. 2 clearly shows non- 2, 4, 5, and 7 constitutes a fairly confined linear trend with a positive Gaussianity: If a joint Gaussian distribution is projected onto a two- slope, whereas Clusters 3, 6, and 8 are plotted below the trend (i.e., on dimensional plane such as Fig. 2, a single cloud consisting of data (a) (d) (d) (g) F_mg- cluster 5 No_mg_L Fig. 2. Major solute concentrations expressed in the logarithmic scale (e.g., log1o (Na mg/L) for x-axis in all diagrams), with discrimination of the eight clusters for a total of 590 groundwater samples. H. Iwamori et al. Journal of Volcanology and Geothermal Research 433 (2023)107739 (a) Na_mg_L (b) K_mg_L Mg_mg_L (d) Ca_mg_L 口 口 cluster 6 cluster 4 6 cluster （f S04_mg_L (h) HC03_mg_L （e) (g) cluster 6 cluster 6 cluster cluster F_mg_L NO3_mg-L G (k) Br_mg-L (1)TotC_cal_mg-L cluste 2 cluster 9 8 2 2 Fig. 3. Mean concentrations in the logarithmic scale of the individual clusters. Small dots of the same color indicate the individual sample values in each cluster. a p p d e pus s appear to be closely associated with volcanoes. lower density as one moves away from the mean. Such non-Gaussian Cluster 3, on the other hand, occurs mainly on the northern margin of characteristics of the data used in this study are key to extract inde- Kyusyu Mountains (in the southern part of study area, Fig. 4) and is pendent features inherited in the data (Iwamori et al., 2017), and the partly distributed in Chikuhi Mountains (in the northern part of study clusters identified by KCA as in Figs. 2 and 3 will be useful to interpret P ns n n the processes and sources that originated the groundwaters. It is noted (Geological Survey of Japan, 2022). Cluster 8, which has the highest that no information concerning geographical distribution of the average concentrations of many dissolved elements, is distributed groundwater samples, isotopic compositions (e.g., oxygen, hydrogen, helium), temperature, or pH, was incorporated into the cluster analysis. Matsumoto, 1993) or the Futagawa-Hinagu fault zone that caused the In the following, how the clusters are related to those parameters of 2016 Kumamoto Earthquake, with the exception of two points at the geographical spatial coordinates, 8180, 8D, 3He/He, temperature, and northwest end of study area. Clusters 1 and 6 are scattered, and espe- pH will be presented. cially Cluster 6 is sparsely found throughout the area (Fig. 4). Now we focus on the Aso caldera area, where Clusters 4, 5, and 7 3.2. Geographical distribution of clusters characteristically occur as stated above, together with a subordinate number of Clusters 1 and 2, as well as a few samples of Cluster 6 and 8 (Fig. 5). The bottom of the Aso caldera (caldera floor) is called Aso-dani Fig. 4 shows the whole studied area, whereas Fig. 5 corresponds to a magnified map around the Aso caldera, both with the sampling locations (Aso valley) in the north and Nango-dani (Nango valley) in the south, of groundwaters and the same color symbols for the eight clusters as in forming a flat topography filled with at least several hundred meters of lake sediments. The northern Aso-dani is widest in the north-south di- Figs. 2 and 3. In Fig. 4, the distribution of each cluster shows charac- rection along the two arrows in Fig. 5, and extends partly to the north of Uchinomaki as if eroding the caldera wall (Fig. 5). both concentrated in two volcanic regions, one in the Aso caldera and the other in the Kuju area, with the exception of one point each in the Cluster 4, the most abundant in the Aso caldera, is distributed on the western foot of Aso and the Uto Peninsula (Fig. 4). Cluster 4 is distrib- caldera floor, and most of the sampling sites are restricted to the lowest part below 500 m in elevation (Fig. 5). However, Cluster 4 is not uted mostly in the Aso caldera, but none in Kuju Volcano. Cluster 2 is distributed in the western part where the caldera wall is absent. Clusters subordinate number in the Aso caldera. The above clusters 2, 4, 5, and 7 1 and 2 spring onto or just below the slope of the caldera wall, H. Iwamori et al. Journal of Volcanology and Geothermal Research 433(2023) 107739 pasec 33.2 Chikuhi Mountains 口 6 33 5 o 4 * 3 + Aso 2 > 32.8 Beppu-Shimabara 14 Plain- Graben L53 32.6 Kyushu Mountains 20km 130.4 130.6 130.8 131 131.2 131.4 131.6 Fig. 4. Geographical distribution of the groundwater clusters in study area. The cluster symbols are the same as in Figs. 2 and 3. The gray scale coding indicates the altitude as in Fig. 1. Several key locations are labeled (see the main text for details). The Beppu-Shimabara graben is indicated by the two pale broad lines, and the Futagawa-Hinagu “Fault Zone" is indicated by a broad green line. The thin solid line represents the coastline (O m above sea level), and the altitude is represented by the same gray color coding as in Fig. 1 (lower is darker, higher is brighter (whiter)). The dashed line outlines the foot of Aso Volcano, bounded by distribution of the Aso pyroclastic flow and topography. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article). lower in this order (Cluster 4, 2, and 1) (Fig. 3), and this distribution- altitude-solute concentration relationship is one of the keys to the elucidation of the causal factors, which will be discussed later. It is noted that, in Fig. 5, Cluster 2 is also distributed outside the caldera in a few 33 8 locations at 400-600 m in elevation, which is similar to the elevation of Cluster 2 inside the caldera. 口 6 On the side of the central cones opposite the caldera wall across the Uchinomak caldera floor, Cluster 7 is mainly distributed (Fig. 5). Cluster 7 is htsuka 3 distributed around the central cones, except in the eastern part of the 2 caldera, resulting in a distribution that traces the shape of the caldera, similar to clusters 4, 2, and 1. In the Uchinomaki area (Fig. 5), located 32.9 Central Cones near the caldera wall on the caldera floor, Cluster 7 hot springs occur exceptionally on the opposite side of the central cones. As already mentioned, the caldera floor is widest in the north-south direction across Uchinomaki. Within this widest area, Cluster 5 is densely located to the south of Hontsuka Volcano which appears as three small peaks con- sisting of felsic lavas, sandwiched between Clusters 4 and 7. 32.8 3.3. Other physicochemical characteristics of clusters 10.km The oxygen and hydrogen isotopic ratios, 8180 and 8D, are useful to 131 131.1 discuss the origin of groundwaters, particularly to distinguish the Fig. 5. Geographical distribution of the groundwater clusters in the Aso meteoric water component and other components such as seawater and caldera. The cluster symbols are the same as in Figs. 2 to 4. The gray scale volcanic fluid components. In Fig. 6, although most of the samples are coding indicates the altitude in meter, based on SRTM15_plus ver.2.1. The two distributed along the meteoric water line, Cluster 8 forms a broad but arrows correspond to the cross section in Fig. 9. The asterisk symbols (*) (uxo Aau Osi yu preo uoap a u pun reau represent the important sites (Nakadake Volcano, Hontsuka Volcano, and Cluster 8 is distributed along the southern edge of the Beppu- Uchinomaki Hotspring). Shimabara graben, a E-W large rift structure of 30 to 40 km wide, and the Futagawa-Hinagu fault zone (Fig. 4). In particular, the easternmost surrounding Cluster 4. Although a few in number, Cluster 2 is distrib- sample is strongly shifted toward a magmatic fluid range, despite its uted between 500 m (just below the caldera wall) and 600 m above sea location (~33.16 N, 131.54 E), in the non-volcanic area on the fore-arc level, and Cluster 1 is distributed between 600 and 900 m. As a result, side of the volcanic front. Together with its high salinity (7810 ppm Na Cluster 4, 2, and 1, in that order, show a concentric zonal distribution, and 11,200 ppm Cl), this sample can be classified as the Arima-type tracing the shape of the caldera. The mean concentrations of solutes are brine (Matsubaya et al., 1973). 5 H. Iwamori et al. Journal of Volcanology and Geothermal Research 433 (2023) 107739 dD_by_CRD dD_by_CRD 8 7 口 6 5 4 3 2 回 0.5 1.5 2.5 d180 Na_mg_L Fig. 6. Left: Oxygen-hydrogen isotopic ratio of the whole samples (x-axis, 818O; y-axis, SD). The solid line represents the annual meteoric water line around Kumamoto (6D =8.5 8180 + 17.7, Kagabu et al., 2011) and the dashed lines indicate directions toward a magmatic water (e.g., 8180 = 7 and D = -32, Kusuda et al., 2014). Right: Na-SD diagram (x-axis, log1o(Na mg/L); y-axis, SD). only for the Aso caldera samples shown in Fig. 5. The cluster symbols are the same as in Figs. 1 and 2. Cluster 6 is also shifted toward high 818o, although the plots of highland plains widely distributed on the caldera rim. Cluster 5 char- Cluster 6 are scattered. Cluster 7 also shows a large shift toward high acteristically occurs around Hontsuka Volcano located on the caldera 8180 for three samples from a hot spring in Kumamoto Plain (~32.89 N, floor and close to the central cones, broadly between the geographical 130.84 E), to the west of Aso Volcano (Fig. 4). Cluster 3 is mainly plotted extents of Cluster 4 and Cluster 7. The mean recharge elevation of along the meteoric water line and is geographically distributed along the Cluster 5 is 922 m, intermediate between Cluster 4 and Cluster 7. basement rocks in the southern mountain range. Several Cluster 3 Cluster 7 on the central cone side has helium isotopic ratios of samples in the western part of the mountain range near the coast are 3He/*He > ~3 Ra where Ra indicates air 3He/*He (Fig. 7) and a high shifted from the meteoric water line toward the seawater composition temperature (average 45.8 °C, Fig. 8), suggesting the influence of vol- 8180 = 0 and 8D = 0. canic gases and heat. On the other hand, Clusters 1 and 2, on the caldera In the Aso caldera, Clusters 6 and 8 rarely appear and no Cluster 3 appears. Groundwaters in the Aso caldera are mostly classified into Cluster 1, 2, 4, 5, and 7, plotted along the meteoric water line. These 3He_4He. Ra oxygen-hydrogen isotope ratios vary systematically in this order (order of increasing dissolved Na concentration), with mean and standard de- viation of SD of -51.65 ± 1.27 %o (Cluster 1), -52.31 ± 1.39 %o (Cluster 2)， -53.14 ± 1.58 (Cluster 4), -54.27 ± 1.83 (Cluster 5), and - 55.59 ± 1.22 (Cluster 7), respectively (Fig. 6 right). In Aso Vol- cano, SD values have been shown to correspond to recharge elevations (Kagabu et al., 2011; Yamada et al., 2011), and following the proposed equation of Kagabu et al. (2011), the above mean 8D values and the standard deviations correspond to the recharge elevations of 762 ± 77 m,802 ± 84 m,853 ± 97 m, 922 ± 112 m, and 1002 ± 74 m, 8 respectively. 7 Most of the highlands on the northern half of the caldera rim are below 950 m in elevation, and elevations above 1000 m are mainly ■ 6 limited to the central cones (and a few peaks on the caldera rim) (Fig. 5). 口 5 Cluster 7, which is distributed on the central cone side, has an average 4 recharge elevation above 10o0 m, suggesting that it is derived from 3 precipitation on the central cones. On the other hand, Clusters 1, 2, and 4 have the recharge elevations <looo m, and show a monotonous 2 variation in this order as follows; as the recharge elevation increases from 762 m (Cluster 1) to 853 m (Cluster 4), the mean Na concentration increases and the elevation at which the groundwater springs out de- Na_mg_L creases from the middle of the caldera wall to the caldera floor (Fig. 5). Cluster 4, the most abundant cluster in the Aso caldera, shows a recharge Fig. 7. He isotopic ratio ?He/*He (normalized to that of air, Ra) as a function of elevation of about 850 m, which is similar to the average elevation of the log1o(Na mg/L) for the eight clusters in the whole study area. The cluster symbols are the same as in Fig. 2. H. Iwamori et al. Journal of Volcanology and Geothermal Research 433(2023) 107739 Temp C Na_mg_L Na_mg_L Fig. 8. Temperature and pH as a function of logio(Na mg/L) for the eight clusters in the whole study area. The cluster symbols are the same as in Fig. 2. wall side, are cooler (around 10 °C) and *He/*He ~ 1 Ra, indicating 4. Discussion negligible volcanic gas or thermal influence. Cluster 4, which is widely distributed on the caldera floor, shows a wide range in ?He/*He from ~1 4.1. Fluid cycling in the Aso caldera Ra to as high as Cluster 7 (blue circles, Fig. 7), and temperatures vary from ~10 °C as low as Clusters 1 and 2 to 40 °C (Fig. 8). Cluster 5, which Based on the characteristics described in Chapter 3, we infer the appears in the caldera associated closely with Clusters 4 and 7, shows an formation processes of Clusters 1, 2, 4, 5, and 7, which occur in the Aso intermediate nature between the two, but has low pH with an average of caldera (Fig. 9). Among them, Cluster 1 springs from the middle to the 5.9 and is the only acidic cluster in the entire study area (Fig. 8). upper part of caldera wall with the lowest solute concentration (Fig. 3), Cluster 3 does not appear in the caldera, but the values of and has a recharge elevation equivalent to the spring altitude. These 3He/*He Ra, temperature, and pH almost overlap Clusters 1 and 2. observations suggest that the meteoric water springs out as Cluster 1 However, a few samples near the southeastern root of the Uto Peninsula without much subsurfaceflow and dissolution of elementsfrom the have a high ?He/*He of about 4.5 Ra and a high Na content (Fig. 7). country rocks (Fig. 9). Cluster 2 has the second lowest solute concen- These samples correspond to the samples in Fig. 6 (left) that are shifted tration (Fig. 3) and springs from the lower part of the caldera wall. The away from the meteoric water line and toward a seawater composition. oxygen-hydrogen isotope ratios lie almost on the meteoric water line, Cluster 6 is characterized by low 3He/*He (Fig. 7), high temperature while the estimated recharge elevation is slightly higher than that of (average 50.5 °C), high pH (average 8.5, Fig. 8), the highest average Cluster 1 (Fig. 6). These features suggest that meteoric water seeped concentration of F in the eight clusters (Fig. 3), and involvement of from the upper part of the caldera wall, reacted with the Aso pyroclastic samples shifted away from the meteoric water line toward high 8180 flow sediments constituting the caldera wall and rim, and flowed down (Fig. 6). Cluster 8, with the highest average salinity, shows relatively incorporating some solutes, and sprang out below the caldera wall. The high ?He/*He (2-5 Ra), high temperature (average 45.8 °C) and meteoric water seeped from the plateau on the caldera rim flowed down neutrality (pH = 7.0), except for one sample. over a longer distance and upwelled to the caldera floor in large Fig. 9. Groundwater cycling model in the Aso ←NNW caldera, illustrating schematically the underground Nakadake (active crater) structures and flow paths with an arbitrary under- ground depth scale. The topography is drawn for the 个 NNW-sSE cross section (see Fig. 5 for the location) Centralcones that passes through Nakadake, Hontsuka Volcano, Caldera rim and Uchinomaki Hotspring, using 5 m-mesh DEM wall floor C7 Altitude (m) 0 (Digital Elevation Model), Geospatial Information ~1000m C4~850m Authority of Japan. The recharge source altitudes in meter for Cluster 1, 2, 4, 5, and 7 (C1, C2, C4, C5, C7) C2 ~800 m are labeled near the corresponding locations, which C1 ~760 m Hontsuka were estimated based on SD (Fig. 6). Uchinomaki 个 500 不 AC5 920m C2 (underground depth is not scaled) ? Basement high-T 33.05 33 32.95 32.9 Latitude (degree N) H.Iwamori et al Journal of Volcanology and Geothermal Research 433 (2023) 107739 Cluster 4 corresponds to this type of groundwater, and shows higher derived from the caldera rim and the central cones). Based on the solute concentrations and recharge elevations than Cluster 2. Ground- knowledge and findings inside the Aso caldera, we discuss the circula- sn t p ann m s tion of groundwaters and their origin outside the caldera in the following sections. and magmatic gases potentially beneath the central cones as illustrated in Fig. 9. 4.2. Origin of groundwaters and deep-seated fuids over study area The influence of magmatic gases and heat sources is most evident in Cluster 7, which surrounds the central cones. Meteoric water infiltrating Cluster 2 is widely distributed around the foot of the vast volume of from the central cone slope at an average elevation of ~1oo0 m is Aso Volcano (Fig. 4). Based on the origin of Cluster 2 inside the Aso caldera (Section 4.1), Cluster 2 groundwaters outside the caldera and high temperatures (Figs. 2, 3, 7, 8). Most of the Cluster 7 water represent meteoric waters percolating through the Aso volcaniclastic samples, which fringe the central cones, have deep spring sources rocks and strata (especially the enormous pyroclastic flow deposits) as 800-1200 m below the surface, and its flow reaches Uchinomaki near partially illustrated in Fig. 9. Kuju Volcano is another active volcano in the caldera wall; the cross-section in Fig. 9 (along the arrows in Fig. 5) study area and is characterized by Cluster 7 and Cluster 5 (Fig. 4), both shows that the caldera floor is widest along this section in the N-S di- of which are related to the volcanic gases, particularly dissolved SO4 as rection, probably because the caldera is deeply gouged (with thicker was discussed for the Aso caldera (Section 4.1). However, Cluster 4, caldera lake sediments). Owing to this structure, flow of Cluster 7 from which predominates on the Aso caldera floor (Fig. 5), is not seen at Kuju the central cone side reaches Uchinomaki without mixing with the Volcano. To produce Cluster 4 type groundwater, a relatively long shallow aquifer (i.e., Cluster 4) through the deep part below the caldera pathway for percolation of groundwater is required to dissolve elements floor, as shown in Fig. 9. from the pyroclastic materials, whereas at Kuju Volcano such long- On the other hand, a part of Cluster 7 groundwater flowing down the distance seepage in pyroclastic material has not occurred, probably relatively shallow depth of the central cone slope turns upward on the due in part to the fact that Kuju Volcano is composed of relatively dense south side of Hontsuka Volcano (right side of Fig. 9) that is mainly lava, such as lava domes, in addition to pyroclastic material. The composed of lava. Then mixing with Cluster 4 groundwater occurs to absence of Cluster 2 in the main body of Kuju Volcano is also consistent produce Cluster 5, which has characteristics intermediate between with a lack of effective infiltration in pyroclastic materials, which is different from Aso Volcano. Cluster 5, estimated from hydrogen isotope ratios, is roughly in the Cluster 8 involves geochemical affinities with the Arima-type brine, middle of 850 m (recharge elevation of Cluster 4) and 1000 m (Cluster and is found along the southern margin of the rift valley or along fault 7), and is consistent with the solute concentrations seen in Fig. 3 being zones. Studies on deep fluids from the Arima region (type locality of the intermediate between the two. However, pH and HCO3 are exceptional: Cluster 5 is acidic with an average pH of 5.9, while Cluster 4 is 7.3 and the Arima-type brine originates from a “slab fluid" derived from a sub- Cluster 7 is neutral at 7.0. Total C in Cluster 5 is intermediate between ducting plate (Philippine Sea Slab) and its dehydration (Kusuda et al., clusters 4 and 7, but has less HCO3 than both, indicating C being dis- 2014; Nakamura et al., 2014; Iwamori et al., 2020). Groundwater clas- solved in Cluster 5 water more as CO2. When considered in conjunction an sr i s s d r s ps with the a high SO4 concentration in Cluster 5 (Fig. 3), this suggests that (Fig. 8) and is also considered to be a deep-seated fluid with a mantle the volcanic SO4 component brought by Cluster 7 dissolved into Cluster component. One possibility is that slab fluids derived from the sub- 5 at lower temperatures and lowered the pH. The relatively high ducting Philippine Sea Plate may have ascended along a large tectonic 3He/*He in Cluster 5 is also consistent with the addition of volcanic line extending deep. Cluster 6 is characterized by the highest fluorine components. On the other hand, Cluster 7, which is also abundant in SO4 content of the eight clusters (Fig. 3), relatively high temperatures and fluids (Figs. 6 and 8), variable ?He/*He from 0.12 to 3 (Fig. 7), and are temperature (Fig. 8). thought to have ascended through varying degrees of reaction with Yamada et al. (2011) investigated the relationship between the crustal rocks (such as granite) at relatively high temperatures in the carbon and oxygen-hydrogen stable isotopic ratios of the groundwaters depths. Cluster 6 is a wide geographical distribution over the study area in the Aso caldera, and found that groundwaters recharged at high (Fig. 4) indicating a pervasive reaction between a widespread flux of altitude has higher stable isotopic ratio of dissolved inorganic carbon deep-seated fluids and crustal rocks. (813Cpic) than groundwater recharged at low altitude. They sugested Cluster 3 occurs mainly along the northern margin of Kyushu that magmatic CO2 mixes into deeper groundwater flowing nearer the Mountains (Fig. 3), and shows characteristics of meteoric water with magma conduit or chamber, and revealed the deep and shallow path- some dissolved components from the rocks (Figs. 2 and 3), possibly ways of groundwaters originated from the central cones in Aso Volcano, permeated through the Mesozoic and Paleozoic basement rocks which is consistent with the model shown in Fig. 9. comprising Kyushu Mountains. However, in the western part of Kyushu Kagabu et al. (2011) showed that the composition of about 100 Mountains near the coast, the groundwaters classified as Cluster 3 show groundwater samples distributed in the Aso caldera can be broadly classified into Ca-HCO3 and Ca-SO4 types, which are supplied from the s caldera wall side and the central cone side, respectively, as in Fig. 9. and 8). In addition, they show high S04 (Fig. 3) and 8180- 8D off the They also suggested that the Ca-HCO3 type has a dissolved ion con- meteoric water line (Fig. 6) indicate an involvement of seawater. These centration that is proportional to the length of the seepage flow path, features suggest mixing of three sources of the groundwater, including which corresponds to Clusters 1, 2, and 4 (C1, C2, and C4 in Fig. 9). In meteoric water, seawater, and mantle He-bearing components (gases or addition, Kagabu et al. (2011) estimated the residence times for several fluids) from depths. Also considering the presence of a hot spring clas- representative samples using groundwater age tracers, tritium and sified as Cluster 7 with high ?He/*He in Kumamoto Plain, deep fluids chlorofluorocarbons (CFCs), with the former flowing for ~20 years and derived from magma or mantle are widely supplied beneath Kumamoto the latter for ~35 years residence time. In this study, by utilizing the Plain, possibly related to the subducted Philippine Sea slab. In summary, Cluster 1 to Cluster 5, and Cluster 7 correspond to newly found that (i) the concentric structures (Fig. 5), (i) occurrence of Cluster 7 (groundwater provided from the central cone side) at Uchi- and/or different lengths of percolation, of which Cluster 7 were fluxed nomaki, and (i) Cluster 5 as a possible mixture of the two types (those remarkably by the volcanic gases. In addition to the shallow water H. Iwamori et al Journal of Volcanology and Geothermal Research 433 (2023) 107739 circulation, we newly found two types of deep-seated fluids, Cluster 6 However, it differs from Fig. ll(a) because it is based only on the and Cluster 8. Considering the geochemical nature, Cluster 8 is likely abundance ratios of the ions projected on these planes, and does not take derived from the subducted Philippine Sea plate and ascends through into account the abundances themselves, nor the multivariate statistical the large tectonic lines and faults without reactions with the crustal properties, including other solutes. As a result, the correspondence of rocks to preserve its original signature (fluid pathway to the surface, the Piper diagram-based clusters (P1-P8) to the KCA-based clusters (C1- expressed by light blue lines labeled “C8" in Fig. 10). Such fluid flux C8) shows significant scatter (Fig. 11(c)); e.g., P7 represented by the from the subducted plate could be pervasive beneath the area, as was orange triangle includes all KCA-based clusters except for C3, while C4 is os uo s (o na oz q pn divided into all Piper diagram-based clusters except for P8. (Fig. 10a) and numerical simulation (Fig. 10b). Such pervasive deep Consequently, the geographic distribution of Piper diagram-based fluids may react with the crustal rocks where the large tectonic lines and clusters obscures the spatial regularity found with KCA-based clusters faults are absent (fluid pathway labeled “C6" in Fig. 10). (Fig. 12), such as the concentric structure in the Aso caldera (Fig. 12(a) and (b)). For example, KCA-based Cluster 4 (blue circle C4 in Fig. 12(a)), which is characteristically distributed at the caldera floor, is subdivided 4.3. Utility of multivariate statistical analysis into five clusters in the Piper diagram-based classification (P1, P2, P3, Fig. ll(a) shows the projection of the KCA clusters onto a cation high SO4 cluster (P4, blue circle) occurs in the norther caldera floor, ternary plot (bottom left), an anion ternary plot (bottom right), and a which does not discriminate C4 and C5 that represents a mixture of the n q d spring waters originated from the caldera rim and the central cone (i.e., clusters are relatively well recognized and discriminated in the cation C5 as a mixture of C4 and C7 in Fig. 9). In addition, KCA-based Cluster 7 surrounding the central cones (orange triangular symbol C7 in Fig. 12 distribution that stretches in the direction of anion in the central (a)) is subdivided into four clusters in the Piper diagram-based classi- rhombus. In the data used in this study, there are only five cations (Na, fication (P5, P6, P7, and P8, Fig. 12(b)). K, Mg, Ca, and Li) out of 12 solutes, which at first glance appears to be given priority in cluster classification in Fig. 11(a), even though the Similarly, geographical provenance identified by the KCA-based cluster distribution in the whole study area (Fig. 4 and Fig. 12(c)) is number of cations is smaller than that of anions. Fig. 2 shows that the variance of anions in each cluster tends to be larger than that of cations not seen in Fig. 12(d); KCA-based Cluster 7 (orange triangle C7 in Fig. 12 (c)), which is distributed almost exclusively in the active volcanic areas (e.g., SO4), resulting in the apparent preferred usage of cations in cluster classification, which is a statistically fair consequence. In other words, of Aso and Kuju and is characterized by high ?He/*He ratios (Fig. 7), does not appear in Fig. 12(d) and is divided into multiple clusters of both the cations and anions were unbiasedly treated in the statistical Piper diagram-based classification. KCA-based Cluster 8, which re- analysis of KCA. sembles Arima-type brines distributed along the tectonic lines (Fig. 4), is Fig. 11(b) shows the Piper diagram, in which Piper (1944) classified also not observed in Fig. 12(d). The most abundant cluster P1 of the the rhombus in the center of Fig. 11(b) based on several triangular re- Piper diagram-based classification (black solid circle, Fig. 12(d)) is gions and their combinations, defining a total of nine composition types. distributed widely over the study area and includes mostly C1, C2, C3, The data used in this study are divided into eight clusters (P1-P8) and and C4 (with minor population of C5, C7, C8). This wide and inclusive are color coded according to the eight triangular regions (Fig. 11(b)), the distribution of P1 also obscures the geographical provenance, such as basis of the original classification of Piper (1944), and their combination discrimination between C2 (red plus symbol) and C3 (green asterisk) in reproduces Piper's original classification (see the caption of Fig. 11). the southern part of study area in Fig. 12(c). Because the classification is based on the abundance ratios of cations These results suggest that multivariate analyses such as KCA are and anions, the clusters are relatively clearly recognized and distin- useful for discussing the origins of groundwater and shows the guished in the cation and anion triangles, respectively (Fig. 11(b)). Kuju Volcano C6 C8 0 20 20 40 40- 60 60 h aa 80 80- luid（vol.%） 100 100- 120 120- 140 140- 3% 0% 3% (b) Slab-derived fluid and melt (a) P-wave velocity (seismic tomography) (numerical simulation>~30km) Fig. 10. (a) Seismic tomography for P-wave velocity along the E-W cross section in central Kyushu (Zhao et al., 2000), and (b) Numerical simulation including subducting plate motion and mantle convection, and the resultant temperature field, fluid generation-migration, and melt generation (modified from Zhao et al., 2000). The spatial distribution of the low P-wave velocity region in (a) and the fluid-melt distribution are broadly similar, suggesting pervasive fluid and melt in the region. In (b), the slab-derived fluid and its migration as porous flow was numerically calculated for the region <~30 km depth, based on Iwamori (1998), whereas the fluid pathways from ~30 km depth to the surface were schematically illustrated. “C8" corresponds to Cluster 8 spring waters, and “C6" corresponds to Cluster 6 waters in Figs. 2 to 8. The “C6" pathway gradually changes its color from light blue (at ~30 km depth) to purple (closer to the surface), which demonstrates the progressive and varying degrees of reaction between the slab-derived fluids and the crustal rocks during the fluid ascent. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article). H. Iwamori et al. Journal of Volcanology and Geothermal Research 433 (2023)107739 (a)KCA-based (b) Piper diagram-based P8 C7 C6 P6 XC5 o td O ￥C3 *P3 +C2 +P2 ·C1 P1 (C) 田 口 口 KCA-based Fig. 11. Comparison of geochemical classification of the groundwaters based on KCA-based and Piper diagram-based clusters. (a) KCA-based clusters, (b) Piper diagram-based clusters, and (c) the correspondence of Piper diagram-based clusters (P1-P8) to KCA-based clusters (C1-C8), where the symbol size indicates the relative number of corresponding samples. In (b), the Piper diagram-based clusters P1 to P8 may reproduce the original nine groups by Piper (1944) as follows. 1: Alkaline earths exceed alkalis (P1 + P2 + P3 + P4 in Fig. 11(c)); 2: Alkalis exceed alkaline earths (P5 + P6 + P7 + P8); 3: Weak acids exceed strong acids (P1 + P2 + P7 + P8); 4: Strong acids exceed weak acids (P3 + P4 + P5 + P6); 5: Magnesium bicarbonate type (P1 + P7); 6: Calcium chloride type (P4); 7: Sodium chloride type (P5 + P6); 8: Sodium bicarbonate type (P8); 9: Mixed type (P3 + P7). advantages over classical Piper diagram-based evaluations. Combina- the eight clusters were broadly categorized into meteoric, volcanic, and tion of the solute-to-solute ratios obtained in a Piper diagram (and/or deep-seated origins. In addition, geographical mapping of the eight cation-anion ternary diagrams) and the individual solute abundances clusters shows systematic relationships between the geochemical fea- based on, for example, a Stiff diagram provides useful and valid con- tures and geographical provenance, including the concentric zonal straints on groundwater systems, as in the previous studies (e.g., the structure of the clusters in the Aso caldera where the meteoric waters groundwater system in the Aso caldera identified by Kagabu et al., percolating through the caldera rim and the volcanic component- 2011). At the same time, the multivariate statistical properties are bearing waters supplied from the central cones encounter and mix at essential for extracting the hidden data structures that are rarely avail- the caldera floor. Upwelling of deep-seated fluids that exhibit able with conventional approaches, and, as demonstrated in this study, geochemical affinities with the Arima-type brine, which could have been are crucial for deciphering the multiple sources and processes involved derived from the subducted Philippine Sea Plate, was identified along in groundwater systems. the tectonic Beppu-Shimabara graben across the central Kyushu and along the Futagawa-Hinagu fault zones. These findings on geographical provenance are not recognized based on the conventional classification 5.Conclusions method of Piper diagram, and suggest that KCA is potentially useful for Multivariate statistical analysis by using whitened data-based k- investigating the origins of groundwaters. means cluster analysis (KCA) has been performed for 12 major solutes (Na, K, Mg, Ca, Li, Cl, SO4, HCO3, F, NO3, Br, and Total C) of a total of Funding 590 water samples in the central Kyushu, SW Japan. Based on the n e ss a This research was supported by the Nuclear Regulation Authority combining the KCA results with other physicochemical data (8180-8D, "Research on Knowledge Development of Giant Eruption Processes, FY 2021" and the Ministry of Education, Culture, Sports, Science and 10 H. Iwamori et al. Journal of Volcanology and Geothermal Research 433 (2023)107739 (a) KCA-based: Aso caldera (b) Piper diagram-based: Aso caldera C8 P8 P7 33 C6 P 32.9 32.8 131 131.1 131 131.1 (c) KCA-based: Whole area 33.2 C8 C7 33 口 C6 C5 o C4 ? C3 32.8 + C2 C1 32.6 (d) Piper diagram-based: Whole area 33.2 P8 P7 P6 P5 P4 32.8 ? P3 + P2 P1 32.6 130.4 130.6 130.8 131 131.2 131.4 131.6 and (c) KCA-based, (d) Piper diagram-based clusters in the whole study area. H. Iwamori et al. Journal of Volcanology and Geothermal Research 433 (2023) 107739 Technology (MEXT) of Japan, under its Earthquake and Volcano Haz- component analyses. Geochem. Geophys. Geosyst. 18, 994-1012. https:/doi.org/ ards Observation and Research Program. 10.1002/2016GC006663. Iwamori, H., Nakamura, H., Chang, Q., Morikawa, N., Haraguchi, S., 2020. Multivariate statistical analyses of rare earth element compositions of spring waters from the Author statement Arima and Ki areas, Southwest Japan. Geochem. J. 54, 165-182. https:/doi.org/ 10.2343/geochemj.2.0583. Kagabu, M., Shimada, J., Shimano, Y., Higuchi, S., Noda, S., 2011. Groundwater flow HI designed the outline of this work. HI, HN, NM, MT, AI, SS sampled system in Aso caldera. J. Japanese Assoc. Hydrol. Sci. 41, 1-17 (in Japanese). and analyzed the samples. HI, HN, SH, TN performed statistical analyses. Kusuda, C., Iwamori, H, Nakamura, H., Kazahaya, K., Morikawa, N, 2014. Arima hot All authors discussed the contents and wrote the manuscript together. 119. https://doi.org/10.1186/1880-5981-66-119. Lopez, D.L, Williams, S.N., 1993. Catastrophic volcanic collapse: Relation to Declaration of Competing Interest hydrothermal processes. Science 260, 1794-1796. Matsubaya, O., Sakai, H., Kusachi, 1., Satake, H., 1973. Hydrogen and oxygen isotopic ratios and major element chemistry of Japanese thermal water systems. Geochem. J. The authors declare that they have no conflict of interest. 7, 123-151. Matsumoto, Y., 1993. Conception of the Beppu-Shimabara graben, its development and Data availability problems. Mem. Geol. Soc. Japan 41, 175-192 (in Japanese). Morikawa, N., Kazahaya, K., Fourre, E., Takahashi, H.A., Jean-Baptiste, P., Ohwada, M., LeGuern, F., Nakama, A., 2008a. Magmatic He distribution around Unzen volcano Data used are attached as Supplementary Material. inferred from intensive investigation of helium isotopes in groundwater. J. Volcanol. Geotherm. Res. 175, 218-230. Acknowledgment Morikawa, N., Kazahaya, K., Masuda, H., Ohwada, M., Nakama, A., Nagao, K., Sumino, H., 2008b. Relationship between geological structure and helium isotopes in deep groundwater from the Osaka Basin: application to deep groundwater The authors would like to thank an anonymous reviewer and Prof. hydrology. Geochem. J. 42, 61-74. Tobias Fischer for their constructive comments. Nakamura, H., Fujita, Y., Nakai, S., Yokoyama, T., Iwamori, H., 2014. Rare earth elements and Sr-Nd-Pb isotopic analyses of the Arima hot spring waters, Southwest Japan: implications for origin of the Arima-type brine. J. Geol. Geosci. 3, 161. Appendix A. Supplementary data https://doi.0rg/10.4172/2329-6755.1000161. Ohwada, M., Kazahaya, K., Itoh, J., Morikawa, N., Takahashi, M., Takahashi, H.A., Inamura, A., Yasuhara, M., Tsukamoto, H, 2012. Passive degassing of magmatic Supplementary data to this article can be found online at https: //doi. volatiles from Iwate volcano, NE Japan, based on three-dimensional measurement of org/10.1016/j.jvolge0res.2022.107739. heliumistpingrouwatephyRe172204htp://r 10.1029/2011JB008532. Piper, A.M., 1944. A graphic procedure in the geochemical interpretation of water- References analyses. Trans. Am. Geophys. Union 25, 914-928. Seno, T., Stein, S., Gripp, A.E., 1993. A model for the motion of the Philippine Sea plate Furukawa, K., Miyoshi, M., Shinmura, T, Shibata, T., Arakawa, Y., 2009. Geology and consistent with NUVEL-1 and geological data. J. Geophys. Res. 98, 17941-17948. petrology of the pre-Aso volcanic rocks distributed in the NW wall of Aso Caldera: Stix, J., Moor, J.M., 2018. Understanding and forecasting phreatic eruptions driven by eruption style and magma plumbing system of the pre-caldera volcanism. J. Geol. magmatic degassing. Earth Planets Space 70, 83. https://doi.org/10.1186/s40623- Soc. Japan 115, 658-671. https://doi.org/10.5575/geosoc.115.658. 018-0855-z. Geological Survey of Japan, AIST, 2022.Seamless digital geological map of Japan V2 1: Takahashi, M., Inamura, A., Takahashi, H.A., Morikawa, N., Togo, Y.S., Kazahaya, K., 200,000. https:/gbank.gsj.jp/seamless/v2.html [Accessed: Mar. 29th, 2022]. Sato, T., Handa, H., Nakama, A., Nakamura, Y., Ohwada, M., Miyakoshi, A., Hurwitz, S., Christiansen, L.B., Hsieh, P.A., 2007. Hydrothermal fluid flow and Tosaki, Y., Tomishima, Y., Daimaru, J., Shimizu, H, Ohsawa, S., Amita, K., deformation in large calderas: Inferences from numerical simulations. J. Geophys. Horiguchi, K., Shibata, T., Koizumi, N., Kawabata, K., Yasuhara, M., 2022. Chemical Res. 112, B02206. https://doi.org/10.1029/2006JB004689. and isotopic compositions of hot spring waters, cold spring waters, groundwaters Hyvarinen, A., Karhunen, J., Oja, E., 2001. Independent Component Analysis. 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tigorite is relatively rich in Fe resulting in low Xmg (Figs. Contents lists available at ScienceDirect Lithos ELSEVIER journal homepage: www.elsevier.com/locate/lithos textural and chemical equilibration with retrograde oliv- ossMark pentinized peridotites of ophiolites: i.e., crystallization geodynamic implications T. Imaoka a*, K. Nakashima b, A. Kamei , T. Itaya d, T. Ohira e, M. Nagashima , N. Kono f, M. Kiji 8 grade metamorphism. In these cases, a set of NiO-Fo Department ofEarthand EnvironmentalSciences,FacultyofScienceYamagataUniversityYamagata990-8560,apan DepartmentfGesciencehmaneiversityatsu90504,an dResearchInstitute ofNtural ciences,kayamaUniversityofScience,1-1Ridai-cho,Kita-ku,kyama000005Jan metamorphic olivine (e.g., Nozaka, 2003, 2011). How- fGakken School,KonoBuilding,27-17Takashima-cho,Numazu410-0056,aan Minoo-higashi High School, 5-4-63 Ao-Gein, Minoo 552-0025,Japan 5b and 5d). The contrastingly high XMg of prograde anti- mate the time necessary for the diffusion. Article history: Cretaceous episodic magmatism produced Nb-rich lamprophyres and adakitic granitoids in the Kinki district of concluding evidence for identification. SW Japan. K-Ar dating of minerals from the lamprophyres, adakites, and hornblende peridotite xenoliths yielded The two types of antigorite have differences in some ages of 109-99 Ma, indicating a short-lived episodic magmatism. The lamprophyres generally display primitive by prograde contact/regional metamorphism (Kunugiza, high-Mg basaltic to basaltic andesite compositions with high Mg# and high Cr and Ni contents that preclude substantial differentiation. Some high-Nb basalt (HNB) and Nb-enriched basalt (NEB) compositions also occur. Keywords: of retrograde antigorite, because it also occurs in Zone II Cretaceous Adakite and variable (La/Yb)n ratios, and can be divided into high-(La/Yb)n (12.5-22.1) and low-(La/Yb)n (3.6-6.1) Lamprophyre serpentinization, Fe in olivine is mainly incorporated into Mantle metasomatism probably resulting from a reequilibration process during Slab rollback dication of increasing chlorite component in antigorite SW Japan NiO from the mantle olivine array (Fig. 4). From this contact aureoles have a slightly magnesian compositions ry thermal anomaly, and hence induce melting of the subducted slab to form adakitic granitoids, and produce not serpentinized peridotite (Ando et al., 2001); much small degrees of partial melting of an enriched metasomatized mantle wedge within the garnet stability field at depths of ≥70 km, whereas the low-(La/Yb)n lamprophyres originated from a different mantle source by a diversity of the Kyoto lamprophyres thus derives primarily from a heterogeneous mantle source that has been variably affected by the results of subduction. ① 2013 Elsevier B.V. All rights reserved. 1. Introduction 2004). Late Jurassic igneous rocks are absent from these regions (e.g., Kiminami and Imaoka, 2013; Kiminami et al., 2009; Sagong et al., Mesozoic igneous rocks and coeval accretionary complexes were 2005), but active magmatism resumed in the Early Cretaceous. Despite developed extensively along the eastern margin of Asia. The Jurassic to Cretaceous accretionary complexes of SW Japan (Fig. 1A, B) indicate ac- magmatism in eastern China was caused by delamination of the litho- tive subduction of an oceanic plate during that time (e.g., Isozaki et al., sphere and/or upwelling of the asthenosphere (e.g, Guo et al., 2006; 1990; Taira et al., 1989). Early to Middle Jurassic igneous rocks are wide- Liu et al., 2009; Tan et al., 2008; Wu et al., 2005a; Xu et al., 2002; Yang ly distributed in the Korean Peninsula and in eastern China (east of the et al., 2007; Zhang et al., 2010). Tan-Lu Fault) (Fig. 1C). Current models suggest that the Korean igneous Magmatism in the Inner Zone of SW Japan began during the latest complexes formed in a continental arc (Kee et al., 2010; Kim et al., 2003, part of the Early Cretaceous (ca. 105 Ma), and corresponding subduc- 2005), whereas the tectonic setting of the eastern China igneous rocks is n n s n still contentious (e.g., Wu et al., 2002, 2005b; Xu et al., 2004; Zhang et al., of SW Japan (Taira et al., 1989). The Inner Zone magmatism is character- guish the ferroan olivine of different origin. In a given * Corresponding author. Tel.: +81 83 933 5765; fax: +81 83 933 5273. known of these magmatic rocks, and hence our understanding of the E-mail address: imaoka@yamaguchi-u.ac.jp (T. Imaoka). subcontinental mantle beneath SW Japan in the Early Cretaceous is 0024-4937/S - see front matter @ 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.lithos.2013.10.014 106 T.Imaokaetal./Lithos184-187（2014)105-131 C Cretaceous plutonic rocks Cretaceous volcanic rocks Jurassic plutonic rocks Jurassic volcanic rocks SanbagawaBelt Shimanto Belt 40°N4 120°E Mino-TambaBelt Chichibu Composite Belt Chugoku BoheiSea Kinki ZTK Kyushu Shikoku Kanto Mts. Fig. 2 Japan Si 125°E RF Pacific Ocean 300km Fig. 1. Index map of Southwest Japan (A), geotectonic divisions of post-Triassic accretionary complexes in the western part of SW Japan (B), and distribution of Jurassic and Cretaceous igneous rocks in and around the Korean Peninsula and the western part of SW Japan (C). Fig. 1c is compiled from Kee et al. (2010), Teraoka and Okumura (2003), and Wu et al. = a :() i s o = o e 's = on = m 's sa = ) s () Line,RB = Ryoke Belt, SyB = San-yo Belt, SiB = San-in Belt. also limited. The earliest Inner Zone magmatism at ca. 105 Ma is charac- geochemical characteristics of the mantle wedge (Ishimaru et al., 2009; terized by adakitic granitoids and high Mg andesites (HMAs) (Kiji et al., Tsuchiya et al., 2005). Studies of these hornblende peridotite xenoliths 1995, 2000), as well as the high-Mg, Nb-rich lamprophyres that are show that the ca. 105 Ma upper mantle beneath the Kinki district had found in the Kinki district (defined as the Kyoto lamprophyre in this been intensely metasomatized (Ishimaru et al., 2009). These 105 Ma rocks can contribute significantly to our understanding of the chemical (Kelemen et al., 2003), we use them for igneous rocks with chemical and physical processes operating in the subcontinental lithosphere be- characteristics that include Si02 ≥ 56 wt.%, Al203 ≥ 15 wt.%, Mg0 neath the district, and to the Early Cretaceous evolution of the eastern usually <3 wt.% and rarely >6 wt.%, low HREE (Yb < 1.9 ppm) and Y margin of the Asian continent. The tectonic setting of the Inner Zone is (<18 ppm), and high Sr (>400 ppm) and Sr/Y (>40), following considered to have changed during the Early Cretaceous, from subduc- n ( e n e n interpreted as indicating both the presence of residual garnet and argued that Cretaceous magmatism in SW Japan was initiated in relation the absence of plagioclase in the source region of the adakites to Early Cretaceous slab rollback in eastern Asia. However, the origin and (Defant and Drummond, 1990; Peacock et al., 1994). In this paper tectonic environment of this initial stage of Cretaceous magmatism, char- we reveal the close space-time association of these adakitic granit- acterized by adakitic granitoids and lamprophyres, and the geodynamic oids and lamprophyre dikes, as well as some hornblende peridotite relationships between this magmatism in SW Japan and the magmatism xenoliths, all of which are dated at ca.105 Ma. Ultramafic xenoliths in the continental area of the Korean Peninsula and eastern China, remain in arc magmas provide direct information on the petrological and unclear. T. Imaoka et al./ Lithos 184-187 (2014)105-131 107 Early Cretaceous; stages II to IV dikes are Late Cretaceous. In this paper raphy, K-Ar ages, whole-rock and mineral chemistry, and Sr-Nd iso- we use the term “Kyoto lamprophyre" for the dikes of stage I, and define topes of the Early Cretaceous igneous rocks from the Kinki district of them as Early Cretaceous lamprophyres within the Mino-Tamba Belt SW Japan, including the adakites, lamprophyres, and peridotite. Our around Kyoto city and its surrounding areas. results provide fresh insights into the causes and geotectonic context The peridotite xenolith described in this paper measures 3 m across tions, the approximate estimation suggests that the cool- and was first found at Ohhiro-dani by Kiji (1987). It occurs within a We also put forward a model that explains how the ca. 105 Ma hornblende spessartite dike in the northern part of Kyoto city. Such xe- magmatism in SW Japan fits into the regional geotectonic evolution of noliths have aroused special interest, because the host rock has been the eastern margin of Asia. dated at 106 Ma, as will be described later. The xenoliths thus provide direct information on the nature of the mantle wedge during the Early 2. Geologic background fonse, B., Zinin, P. and Sharma, S.K. (2009) Onset and pro- and geochemical characteristics of the xenolith, and inferred certain 2.1. Jurassic-Cretaceous accretionary complexes in SW Japan especially hydration due to the invasion of hydrous magma. Southwest Japan is divided by the Median Tectonic Line (MTL) into The Tamba granitoids form many small stocks in northern Kyoto the Inner and Outer zones (Fig. 1A). The Inner Zone is generally under- city. The occurrence, petrography, and chemistry of these stocks have lain by a subduction-accretionary complex of Late Triassic to earliest already been described by Kiji et al. (2000), so only their general Cretaceous age (Nakae, 2000; Shuto and Otsuka, 2004; Takami and within the serpentinized mantle wedge following rapid Itaya, 1996) called the Mino-Tamba Belt (Fig. 1B). This belt consists of quartz diorites, and quartz monzodiorites, and they belong to both clastic rocks and oceanic materials such as greenstone, chert, and lime- the ilmenite-series and magnetite-series. They have adakitic character- stone. The geochemistry of the greenstones suggests origin in oceanic istics, and have higher Sr/Y ratios and lower Y contents than the sur- island and plateau settings (e.g., Ichiyama et al., 2008; Sano, 1989). rounding granitoids of the San-yo and Ryoke belts. Hornblende and The Outer Zone of SW Japan is mainly underlain by the following magnitude (Dohmen et al., 2007). Actually, the difference Jurassic-Cretaceous subduction-accretionary complexes: the Sanbagawa adakitic granitoids (Table 2). These ages are older than those of geologic time frame. Although exact time required for granitoids in the San-yo and Ryoke belts (Kiji et al., 1995). The Tamba even if intense low-T serpentinization has obscured the granitoids are thus regarded as the forerunners of the Cretaceous and metabasites, and radiometric ages indicate metamorphism in magmatism of SW Japan. the Cretaceous (Itaya et al., 2011); most of the belt represents the meta- 2007b). Using the latter D in this study, the time was 3. Samples and analytical techniques exhumation immediately after the amphibolite-facies during the Jurassic and Early Cretaceous, and consists of clastic rocks Samples were collected in and around Kyoto and Uji cities in the and oceanic materials such as cherts, limestones, and greenstones of Kinki district. The freshest samples were examined optically, and they lites. Approximate estimation of time required for the ob- include 57 samples of Kyoto lamprophyres and other dike rocks, 3 sam- Shimanto Belt was accreted during the Cretaceous-Early Miocene, and ples of hornblende peridotite xenolith, and 56 samples of the Tamba it is composed of clastic rocks with subordinate amounts of oceanic granitoids. material. Major and trace elements were analyzed using a Rigaku RIX 3000 XRF spectrometer at the Center for Instrumental Analysis of Yamaguchi 2.2. Cretaceous igneous rocks in the Kinki district, inner zone of SW Japan SUPPLEMENTARYMATERIALS and errors and analytical precision of the XRF analyses have been The Cretaceous to Paleogene igneous rocks in the Inner Zone show a documented by Umemoto et al. (2000). Loss on ignition (LO1) was de- zonal distribution parallel to the arc, and have been subdivided into the termined after ignition at 1000 °C for 1 h. CIPW norm calculations Ryoke, San-yo, and San-in belts, from the MTL towards the Sea of Japan were made using the program SINCLAS (Verma et al, 2002). Analyses (Fig. 1C). The Ryoke Belt is characterized by Cretaceous plutono- of rare earth elements (REEs) and additional trace elements in repre- metamorphism, and is associated with the Ryoke low-P/T type regional sentative subsets of 35 samples were made using a Perkin-Elmer tain the minimum duration of cooling during or after Optima 3000 ICP-MS at Activation Laboratories of Ancaster, Canada unfoliated granitoids. The San-yo Belt is characterized by Cretaceous (http:/www.actlabs.com/), employing lithium metaborate/tetraborate magmatism that produced voluminous felsic pyroclastic rocks and coe- fusions. Errors and the analytical precision of the ICP-MS analyses are val granitoids. The San-in Belt is also characterized by felsic to interme- presented on the Actlabs web site. diate volcanic rocks and coeval shallow intrusive rocks, although most Mineral chemical compositions were determined using JEOL JXA- radiometric ages in the belt are Paleogene. spectrometer in his laboratory at Okayama University. Various dikes intrude the Mino-Tamba Belt in the Kinki district of netite inclusion, and having a composition richer in Fe ~ 4.0 × 107 years at 500 °C and 0.5 GPa. An unreason- nization of the Oeyama ophiolite (Nozaka, 2014a). The 15 kV, a specimen current of 20 nA, and a beam diameter of 1-5 μm. characteristics, and rare cross-cutting relationships. From oldest to Wavelength-dispersion spectra were collected with LiF, PET, and TAP pentinization in the Oeyama ophiolite. Intracrystalline petrophysics of serpentinites from MAR 15°N (ODP Leg study); II, hornblende clinopyroxene spessartites (95-90 Ma; for background measurements. The ZAF method was used for data Kutsukake et al., 2010); Ill, clinopyroxene hornblende porphyries correction. Under the conditions described, analytical errors are ±2% (89-82 Ma; this study), and granite porphyries and felsites for major elements and ± 5% for trace elements, as estimated from the (95-80 Ma; Sawada et al., 1994); and IV, andesites (74.8 ± 1.9 Ma; reproducibility observed in multiple measurements. Kiji, 2005) and mafic dikes (67.4 ± 1.7 Ma; Kiji and Kitani, 2009). K-Ar age determinations were carried out on hornblende separates The dikes of stages I and Il are easily distinguished, because stage I dikes are dominated by hornblende and plagioclase with <2 vol.% tites. The method of sample preparation for K-Ar dating has been docu- period in the region of serpentinized mantle wedge fol- mented in detail by Imaoka et al. (2011). Purity of minerals in the clinopyroxene phenocrysts (5%-13%). Stage IV dikes are related to the fractions was estimated to be >99%. Potassium was analyzed using magmatism that produced the surrounding large-scale rhyolites and flame photometry, and argon was analyzed by isotope dilution methods granites of the San-yo Belt. In terms of ages, only the stage I dikes are using a 38Ar spike on a 15-cm-radius sector-type mass spectrometer 108 T. Imooka et al. / Lithos 184187 (2014) 105131 Pliocene-Quaternary 135°30'E Felsicdike Kobe Group Porphyry Granitoidsin theSanyoBelt Lamprophy ★ Tamba granitoids in the Sanyo Belt X Granitoidsin theRyokeBelt Rhyolitic pyroclasticrocks Sasayama Group Tamba Group (incl.Takatsuki F.) Sasayama aka et al., 1982) with a single collector system at Okayama University of Science, 4. Petrography and mineralogy Japan. The analytical procedures, accuracies, and reproducibilities of the methods of K and Ar analysis have been described in detail 4.1. Kyoto lamprophyre by Itaya et al. (1991) and Nagao et al. (1984). The KAr ages were The Kyoto lamprophyres are compact and dark green to black 入β = 4.962 × 10-10/y, and 40 K/K = 0.0001167 (Steiger and Jager, Amphibole crystals are visible to the naked eye. Under the micro- 1977). Age errors are quoted at two standard deviations (2o). scope, specimens show panidiomorphicgranular textures and/or Extractions of Sr and Nd from the powdered whole-rock and holocrystalline porphyritic textures. Some samples can be called mineral samples were carried out at Yamaguchi University, follow- appinite (Fig. 3A: coarse-grained equivalents of closely associated ing the methods of Kagami et al. (1987). Isotope measurements hornblenditic lamprophyres, after Rock, 1991). Major constituent were made at the Department of Geoscience, Shimane University. minerals are plagioclase (650 vol.%) and amphibole (34%69%). Japan, using a Finigan MAT-262 mass spectrometer equipped Plagioclase phenocrysts (0.12.0 mm; An = 0%27%) are euhedral, with 5 Faraday cups. The analytical methods and the accuracy of and in some samples the plagioclase in the groundmass is feather-like. the isotope data have been described by lizumi et al. (1994, 1995) More than half the examined samples contain round pseudomorphs of and lizumi (1996). Measured 87Sr/86Sr and 143Nd/144Nd ratios were nor- olivine (<0.5 vol.%) that are completely replaced by chlorite, serpentine, ads 610 = PN/PN pue 6110 = JS/S 0 pae and other clay minerals; the olivine pseudomorphs are surrounded The 87Sr/86Sr ratio of the NIST SRM987 Sr standard and the 143Nd/144Nd ratio of the JNdi-1 Nd standard were measured during this study. The the olivine. Fine-grained alkali feldspar (Or = 88%98%) is present in respective mean ratios were 0.710236 ± 0.000008 (2o, n = 12) and minor quantities between the plagioclase laths. PN/PNe pue JSgg/ISs panseaW (9 =u) S000000 L60Z1S0 Hornblende phenocrysts (0.11.9 mm) are euhedral to subhedral ratios were corrected relative to reference values of 0.710242 (Kagami with a sieve-like appearance, and show zoning from light brown cores et al., 1989) and 0.512115 (Tanaka et al, 2000), respectively. We used to light green rims. They are often hollow and are sometimes opacitized. the following Chondritic Uniform Reservoir values (= O Ma) for calculat- The amphiboles are rich in TiO2 (up to 4.1 wt.%) and AlOs (up to 2800 = JS/ 00 =JS/S sane (1)N3 pue (1)s3 Su 13.4%), and Mg# values [Mg/(Mg + Fe2+)] exhibit a wide range 143Nd/144Nd = 0.512638, and 147Sm/44Nd = 0.1966 (Goldstein et al, from 0.40 to 0.89 (Table 1A). According to the classification of 1984). Leake et al. (1997). they plot in the fields of magnesiohastingsite, T.Imaoka et al./Lithos 184-187(2014)105-131 109 A B 0.51 ID R Bt TC 0.1mm Fig. 3.(A) and (B) Photomicrographs of Kyoto lamprophyres (spessartites).(A) Coarse-grained spessartite (appinite). (B) Pseudomorph of an olivine xenocryst surrounded by amphibole grains. (C) Hornblende peridotite xenolith in the Kyoto lamprophyre (spessartite). (D) Sagenitic texture of the Tamba granitoid.Biotite contains slender, needle-like rutiles 0.1-2.0 μm in Tlc = talc, Bt = biotite, and Rt = rutile. pargasite, edenite, magnesiohornblende, ferrohornblende, tschermakite, (Or = 79%-97%) are present in minor quantities between the plagioclase ferrotschermakite, and actinolite (Fig. 4A, B). The dated hornblende laths as anhedral grains. (KY-44) can be classified as magnesiohastingsite, based on its bulk The amphiboles are euhedral to anhedral grains up to 1 cm; they are chemical composition; it has high contents of HFSEs (TiO2 3.37 wt.%, pleochroic with brownish green to light green color parallel to the Z Nb 39.1 ppm, Zr 206 ppm, Ta 1.87 ppm, Hf 6.3 ppm), as listed in optic direction, and they are colorless parallel to X. All are calcic amphi- Appendix 1. boles (Fig. 4A; classification of Leake et al., 1997) with Mg# values of Clinopyroxene (<1.8 vol.%) occurs as aggregates of anhedral grains 0.60 to 0.81, and they can be classified as magnesiohastingsite, that measure 0.05 to 0.5 mm; the clinopyroxene may be xenocrystic. pargasite, edenite, magnesiohornblende, and actinolite (Table 1). Mg# values are 0.72-0.90 (Table 1B), and the clinopyroxenes are Biotites are reddish-brown parallel to the Z optic direction, and they moderately enriched in TiO2, Al2O3, Cr2O3, and Na20 (up to 1.48, 5.25, have high Mg# values [Mg/(Mg + Fe)] of 0.48-0.64, and contain greater 1.10, and 0.71 wt.%, respectively) relative to those in the Late Creta- TiO2 contents (2.8-5.0 wt.%) than biotites in the Late Cretaceous granitoids ceous volcanic rocks of SW Japan (Fig. 5; Imaoka and Murakami, 1979). of SW Japan (Fig. 6; Czamanske et al., 1981; Honma, 1974; Kanisawa, Apatite is the most common accessory mineral, existing as laths and 1979; Murakami, 1981; Tainosho, 1973, 1986; Tainosho et al., 1979). The needles (up to 1 mm long) that are commonly full of tiny inclusions. biotites occasionally contain slender, rutile needles 0.1-2.0 μm in width Cr-spinel occurs in close association with the olivine, and it is Cr-rich (most are <0.5 μm) intersecting at an angle of 60° forming equilateral tri- with Cr/(Cr + Al + Fe) ratios of 0.545-0.786, and Al/(Cr + Al + Fe) angle and asterisk-shaped units; such a pattern has been referred to as ratios of 0.141-0.277. Anhedral grains of magnetite, ilmenite, and pyr- sagenitic texture (Fig. 3d). The needles forming asterisks are generally rhotite occur in accessory or minor amounts. The matrix (6-31 vol.%) shorter than those forming equilateral triangles, but both have the same is composed mainly of plagioclase laths with a grain size up to origin. The rutile needles contain up to 0.27 wt.% Nb2Os (Table 1). 0.2 mm. The alteration observed in some specimens was probably the Other accessory minerals include titanite, pyrrhotite, pyrite, ilmen- result of their high volatile content. Secondary minerals include calcite, ite, magnetite (±), chalcopyrite (±), apatite, allanite, epidote, and chlorite, epidote, and pyrite. Some samples contain xenocrysts of quartz zircon. Titanite (up to 1 to 5 mm) is abundant in some samples (up to (0.03-0.5 mm) that show complex reaction relationships with the 2.9 vol.%); these contain minor amounts of Al2O3 and MgO, but no de- surrounding matrix. tectable Nb2Os (Table 1). Ilmenites (0.01-0.2 mm) are rich in hematite molecules (10-62 mol%), and moderately rich in pyrophanite (2.2- 4.2. Tamba granitoids 8.8 mol%). Magnetites (0.01-1 mm in size) are almost pure but with some ulvospinel contents (0-4.1 mol%). The Tamba granitoids are tonalites, quartz diorites and quartz monzodiorites with color index values of 9.0% to 44.4%. Major constituent 4.3. Hornblende peridotite xenolith minerals are plagioclase (47-67 vol.%), K-feldspar (0.1-8.5 vol.%), quartz (6-23 vol.%), amphibole (3.5-38.9 vol.%), and biotite (4.7-14.8 vol.%) The peridotite xenolith has been highly metasomatized, and is mainly (Kiji et al., 2000). Plagioclase phenocrysts (1-5 mm; An = 11%-38%) composed of olivine (43 vol.%) and hornblende (34 vol.%) (Fig. 3B). The are euhedral to subhedral. In some porphyritic samples the plagioclase rest of the rock consists of plagioclase, clinopyroxene (<2.1%), Cr-spinel, occurs as large phenocrysts up to 8 mm in length. Quartz is anhedral, orthopyroxene, and apatite. Secondary minerals include talc, chlorite, and occurs as interstitial grains up to 6 mm in diameter. Alkali feldspars carbonates, and sulfides. Sulfide globules up to 3 cm are sometimes 110 Table 1 Representative electron microprobe analyses of (A) amphiboles, and (B) clinopyroxenes, Cr-spinels, biotite, ilmenite, magnetite, titanite, and rutile from the Kyoto lamprophyres, hornblende peridotite xenoliths, and the Tamba granitoids. Rock type Spessartite Hornblendeperidotitexenolith Adakiticgranitoids Sample no. KY-11 KY-33 KY-11 KY-11 KY-04 KY-11 KY-11 KY-11 KY-33 KY-02X KY-02X KY-02X KY-02X KY-02X KTS-4 KTS-2 TG-03 Point no. Brown core Pale brown Brown Green cOlorlessrim Magnesio- Magnesio- Tschermakite Titanian Titanian Actinolite Ferro- Ferro- Edenite Magnesio- Magnesio- Magnesio- Magnesio- Ferro- Magnesio- Actinolite Magnesio- hastingsite hastingsite pargasite pargasite hornblende actinolite hastingsite hornblende hornblende hornblende hornblende hastingsite hornblende SiO2 41.75 43.54 42.37 42.21 41.18 55.78 47.95 50.26 44.53 41.55 47.06 45.68 51.66 51.52 42.73 53.77 45.85 TiO2 3.04 1.11 2.51 2.51 4.05 0.05 0.21 0.06 0.98 2.13 0.74 0.79 0.18 0.10 2.90 0.19 1.41 13.28 11.37 13.09 12.91 5.30 4.06 14.52 1.37 11.81 Al2O3 13.40 10.77 9.90 10.08 3.79 2.11 1.77 Cr203 0.00 n.d. 0.07 0.02 0.00 0.16 0.02 0.12 n.d. 0.02 0.14 0.01 0.00 0.01 0.14 0.00 000 Feo* 10.75 16.08 12.71 10.78 11.54 11.00 22.71 21.72 17.60 11.48 11.14 15.63 18.23 23.33 9.37 10.87 14.99 Mno 0.13 0.37 0.24 0.14 0.21 0.44 2.37 1.88 0.42 0.18 0.25 0.38 1.76 4.07 0.10 0.79 0.3 et Nio n.d. 0.00 0.01 0.02 0.00 0.00 0.03 0.00 0.00 0.00 0.00 0.00 0.00 0.03 0.04 n.d. 0.00 MgO 13.79 11.74 13.84 13.66 13.26 16.15 8.69 9.65 11.11 13.80 15.25 12.49 10.88 6.59 15.47 16.99 12.66 /Lithos Cao 11.33 10.97 11.83 12.01 12.17 12.67 9.28 9.19 11.14 11.87 11.49 11.20 11.75 11.39 11.63 12.35 10.91 Na20 1.79 1.91 1.73 1.65 1.98 0.18 0.80 0.50 1.93 2.22 1.65 1.57 0.39 0.14 2.26 0.38 2.10 K20 0.89 0.62 0.83 0.82 0.71 0.03 0.31 0.18 0.54 0.58 0.47 0.47 0.12 0.07 1.17 0.15 0.79 F 0.14 0.14 0.07 0.21 0.00 0.12 0.10 0.06 0.06 0.22 0.18 0.08 0.00 0.00 0.23 Cl 0.02 0.08 0.02 0.01 0.02 0.00 0.05 0.03 0.11 0.02 0.07 0.09 0.01 0.00 0.13 (2014) Ex.0= 0.06 0.08 00 0.09 0.00 0.05 0.05 0.03 0.05 0.10 60'0 0.05 0.00 0.00 0.13 96.85 97.86 99.28 96.85 98.53 98.31 97.75 97.71 99.13 98.50 98.26 98.43 98.80 98.59 98.59 98.59 98.59 105- 0=23 WS 6.113 6.415 6.082 6.196 5.996 7.907 7.259 7.551 6.519 5.999 6.722 6.646 7.578 7.853 6.217 7.676 6.682 2.004 2.001 1.887 1.585 1.918 1.804 600 0.741 0.449 1.481 1.278 1.354 0.422 0.147 1.783 0.324 1.318 Sum T 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 8.000 AM 0.404 0.389 0.296 0.430 0.295 0.202 0.203 0.271 0.378 0.469 0.389 0.375 0.233 0.099 0.243 0.031 0.285 0.334 0.123 0.271 0.277 0.444 0.005 0.024 0.006 0.107 0.231 0.079 0.086 0.020 0.011 0.318 0.020 0.154 0.000 0.000 0.007 0.002 0.000 0.018 0.002 0.014 0.000 0.003 0.015 0.002 0.000 0.001 0.016 0.000 0.000 0.424 0.586 0.638 0.342 0.238 0.019 0.453 0.214 0.512 0.518 0.514 0.545 0.148 0.033 0.288 0.226 0.466 3.010 2.579 2.961 2.990 2.877 3.413 1.961 2.160 2.425 2.969 3.248 2.709 2.379 1.498 3.356 3.615 2.751 0.828 1.323 0.826 0.959 1.146 1.285 2.357 2.335 1.578 0.810 0.755 1.283 2.087 2.941 0.779 1.072 1.344 Mn 0.000 0.000 0.000 0.000 0.000 0.053 0.000 0.000 0.000 0.000 0.000 0.000 0.133 0.416 0.000 0.036 0.000 5.000 5.000 4.999 5.000 5.000 4.995 5.000 5.000 5.000 5.000 5.000 5.000 5.000 4.999 5.000 5.000 5.000 0.064 0.072 0.000 0.007 0.020 0.000 0.064 0.181 0.065 0.058 0.061 0.073 0.000 0.000 0.073 0.000 0.017 Mn 0.016 0.046 0.030 0.017 0.026 0.000 0.303 0.240 0.052 0.022 0.031 0.047 0.086 0.110 0.012 0.060 0.037 Ni 0.000 0.000 0.000 0.000 0.001 0.000 0.004 0.004 0.000 0.003 0.000 0.000 0.003 0.000 0.000 0.000 0.000 Ca 1.777 1.731 1.820 1.889 1.899 1.925 1.505 1.480 1.747 1.836 1.759 1.746 1.847 1.860 1.813 1.889 1.704 Na 0.143 0.151 0.150 0.087 0.054 0.051 0.124 0.095 0.136 0.081 0.149 0.134 0.064 0.030 0.102 0.051 0.242 Sum B 2.000 2.000 2.000 2.000 2.000 1.976 2.000 2.000 2.000 2.000 2.000 2.000 2.000 2.000 2.000 2.000 2.000 Na 0.365 0.396 0.332 0.382 0.504 0.000 0.112 0.049 0.412 0.539 0.308 0.310 0.049 0.011 0.536 0.054 0.352 K 0.167 0.117 0.152 0.154 0.132 0.006 0.059 0.035 0.101 0.107 0.085 0.087 0.022 0.014 0.216 0.027 0.147 sum A 0.532 0.513 0.484 0.536 0.636 0.006 0.171 0.084 0.513 0.646 0.393 0.397 0.071 0.025 0.752 0.081 0.499 Total 15.532 15.513 15.483 15.536 15.636 14.977 15.171 15.084 15.513 15.646 15.393 15.397 15.071 15.024 15.752 15.081 15.499 0.065 0.066 0.032 0.098 0.000 0.054 0.048 0.029 0.028 0.102 0.082 0.037 0.000 0.000 0.106 0.005 0.020 0.005 0.003 0.005 0.000 0.013 0.008 0.028 0.005 0.017 0.022 0.002 0.000 0.032 Mg# 0.77 0.65 0.77 0.75 0.71 0.73 0.45 0.46 0.60 0.77 0.80 0.67 0.53 0.34 0.80 0.77 0.67 Rock type Spessartite Peridotite xenolith Adakitic granitoids Sample no. KY-04 (11) KY-04 (38) KY-04 (14) KY-04 (1) KY-04 (2) KY-02X KY-51 KY-48 KY-02X KY-02X KY-02X TG-03 HA-03 TAW-01 TAW-01 TAW-01 TAW-01 Mineral xd xd Cpx Cr-spinel Cr-spinel xd xd xd xd Cr-spinel Cr-spinel Biotite Biotite Ilmenite Magnetite Titanite Rutile SiO2 53.38 54.55 49.77 0.00 0.00 53.31 53.52 51.56 51.40 0.05 0.00 38.43 37.17 0.00 0.00 33.74 1.19 TiO2 0.25 0.24 1.07 0.94 0.15 0.17 0.30 0.61 0.51 0.33 0.32 4.27 4.97 47.11 0.06 22.56 97.43 Al2O3 2.40 1.37 5.25 13.12 7.04 1.60 2.64 3.93 2.89 14.35 10.38 14.03 13.48 0.04 0.04 10.99 0.26 Cr203 1.01 0.57 60'0 43.87 56.72 46.78 54.65 0.00 0.12 0.00 00 Feo* 3.37 4.49 7.29 32.38 27.74 2.56 3.94 6.44 7.15 28.20 20.60 16.33 19.68 49.47 92.80 2.67 0.43 Mno 0.12 0.16 0.20 0.38 0.39 0.06 0.14 0.20 0.10 0.47 0.34 0.22 0.25 1.95 00 0.07 0.05 Nio 0.05 0.02 0.01 0.08 0.01 0.14 0.09 0.00 0.00 0.02 0.03 MgO 17.00 17.79 13.77 8.76 7.18 17.41 16.87 15.75 15.53 8.69 12.11 12.59 11.36 0.05 0.00 3.19 0.00 21.97 0.00 0.00 Cao 22.09 20.33 21.67 0.15 0.16 23.10 22.36 21.65 0.02 0.00 0.00 0.00 24.12 0.43 Bao 一 一 0.13 Na20 0.28 0.20 0.23 0.27 0.28 0.26 0.20 0.13 0.11 0.29 K20 0.01 0.01 0.00 0.02 0.01 0.03 0.00 9.68 10.12 0.25 V205 0.18 0.09 0.12 0.06 Nb205 0.30 F 一 1.14 0.03 Ex.0= 0.49 Total 99.96 99.73 99.35 99.86 99.48 97.37 100.80 101.14 99.43 99.15 98.55 95.68 97.26 98.64 92.99 98.69 99.79 T.Imaoka et al. 0=6 0=6 0=6 0=32 0=32 0=6 0=6 0=6 0=6 0=32 0=32 0=22 0=22 0=6 0=32 0=5 0=2 !S 1.943 1.984 1.857 0.000 0.000 1.979 1.937 1.883 1.912 0.013 0.000 5.751 5.603 0.000 0.000 1.081 0.016 Ti 0.007 0.007 0.030 0.185 0.031 0.005 0.008 0.017 0.014 0.065 0.063 0.481 0.563 1.804 0.014 0.544 0.978 Al 0.103 0.059 0.231 4.048 2.270 0.070 0.113 0.169 0.127 4.443 3.213 2.475 2.395 0.002 0.015 0.415 0.004 ΛD 0.029 0.016 0.003 9.079 12.267 9.716 11.349 0.000 0.014 0.000 0.007 0.000 0.000 /Lithos 0.038 0.020 0.000 0.000 0.025 0.013 一 Fe3+ 2.465 1.382 1.660 1.298 0.390 15.95 0.103 Fe2+ 0.137 0.228 4.623 4.964 0.079 0.119 0.197 0.222 4.535 3.227 2.044 2.481 1.716 7.999 0.072 0.005 184-187 Mn 0.004 0.005 0.006 0.084 0.090 0.002 0.004 0.006 0.003 0.105 0.076 0.028 0.032 0.084 0.008 0.002 0.001 Ni 0.001 0.001 0.000 0.017 0.002 0.030 0.019 0.000 0.000 0.001 0.007 0.923 0.965 0.766 3.418 2.928 0.963 0.910 0.858 0.861 3.403 4.742 2.808 2.553 0.004 0.000 0.152 7(2014)105- 0.862 0.792 0.866 0.042 0.047 0.874 0.896 0.875 0.863 0.006 0.000 0.000 0.000 0.000 0.000 0.828 Ba 0.002 0.020 0.014 0.017 一 0.019 0.020 0.018 0.014 0.038 0.032 Na 0.018 一 K 0.000 0.000 0.000 0.001 0.000 0.001 0.000 1.847 1.946 0.010 -131 Nb Total 3.995 3.980 4.004 23.999 24.001 3.992 4.007 4.024 4.016 24.001 24.000 15.472 15.619 4.001 24.000 3.124 1.004 F 0.116 C1 一 一 一一 0.002 0.900 0.876 0.771 0.924 Mg# 0.884 0.813 0.795 0.575 0.504 0.679 # 0.692 0.844 一 0.686 0.779 Mg# = Mg/(Mg + Fe), Cr# = Cr/(Cr + Al) 三 112 T.Imaoka et al./Lithos 184-187(2014)105-131 A (Na + K)A ≥ 0.50 B (Na + K)a < 0.50 tremolite magnesiohornblende tschermakite edenite pargasite . actinolite 8 magnesio- hastingsite 4 0.5 0.5 ferropargasite V ferro- hastingsite actinolite ferro-edenite ferrohornblende ferrotschermakite 7.5 6.5 6 5.5 8 7.5 7 6.5 6 5.5 Si in formula Si in formula C D 15 TiO2 (wt.%) o Kyoto lamprophyres AlO3 (wt.%) : · peridotite xenoliths v Late-K VR 10- 品 8 °。 V。 ： 0 0- 0.8 0.6 0.4 0.2 0.8 0.6 0.4 0.2 Mg# Mg# KVR).Amphibole terminology is from Leake et al. (1997). (C) Mg# vs. TiO2 wt.% in the amphiboles. (D) Mg# vs. Al2O3 wt.% in the amphiboles. observed; these are a mixture of pyrrhotite and very small amounts of core to rim (Ishimaru et al., 2009). The Cr# [=Cr/(Cr + Al) atomic chalcopyrite and pyrite. Details of the petrography and mineral chemistry ratio] of chromian spinel varies from 0.66 to 0.88, and there is a wide have been documented by Ishimaru et al. (2009). range of Mg# (0.28-0.70). A positive correlation (r = 0.962) exists be- Olivines have been partly or completely altered to serpentine, chlo- tween the Fo contents of olivines and the Cr# of the chromian spinels, rite, and talc. The olivines exhibit a wide range of Fo from 94 to 87 consistent with the trends (Arai, 1994) observed in high-Mg and -Cr (=100 × Mg#), and their Nio contents decrease from 0.51 to magmas (Ishimaru et al., 2009). 0.22 wt.% with decreasing Fo. The Fo and Nio contents seem to be The amphiboles in the peridotite xenolith are euhedral to anhedral, lowered by contact with hornblende, but grains not in contact with and are either interstitial to the olivine, or crosscut or enclose olivine that phase show only a slight decrease in Fo content (94 to 93) from grains. They are mainly calcic amphiboles (Fig. 4A; classification of 1.8 TiO2 (wt.%) oKyoto lamprophyres 6 Al2O3 (wt.%) Na2O (wt.%) 1.6- · peridotite xenoliths 5- o 。 1.4- 0.8- Late-KVR 1.2 。 8 4 1 0.6- 3 0.8- 0.4 0.6- 0.4- 0.2 0.2- 。 0 0.90.80.7 0.60.50.40.3 0.90.80.70.6 0.50.4 0.3 0.3 Mg# Mg# Mg# Fig. 5. Mg# vs. TiO2, AlO3, and NazO in clinopyroxenes from the Kyoto lamprophyres (spessartites） and peridotite xenoliths. Clinopyroxenes from the Late Cretaceous volcanic rocks (Late-K VR) in western Chugoku district (Imaoka and Murakami, 1979) are also plotted for comparison. T.Imaoka et al./Lithos 184-187(2014)105-131 113 Colorless tremolitic and actinolitic amphiboles attached to the edges of the brown zoned hornblendes often occur as late-stage fibrous or needle- shaped aggregates. Amphiboles with the most prominent variations show decreases in Mg# from 0.80 (brown core) to 0.34 (colorless rim), and decreases in TiO2 from 2.13 to 0.10 wt.% (Table 1A). There is also a decrease in (Na20 + K20) from 2.80 wt.% (core) to 0.21 wt.% (rim). MnO increases from 0.18 wt.% (brown core) through 1.76 wt.% (green) to 4.07 wt.% (colorless rim). Tremolitic and actinolitic amphiboles have low contents of TiO2 (<0.10 wt.%), Al203 (<0.60 wt.%), and Na20 + K20 (<0.05 wt.%), and high values of Mg# (0.81 to 0.92) (Fig. 4B-D). F and !. + Tamba granitoids Cl concentrations are very low in all the amphiboles (<0.22 wt.%). Late-K granitoids Clinopyroxenes have Mg# values of 0.79-0.92, and are moder- ately rich in TiO2 (<0.61 wt.%), Al2O3 (<4.1 wt.%), and Na2O 10 20 30 40 50 60 70 (<0.45 wt.%), and these compositions are nearly identical to those of Mg# value [Mg/(Mg+Fe)] the clinopyroxenes in the spessartites (Fig. 5). It follows, therefore, that Fig. 6. Mg# vs. TiO2 in biotites from the Tamba granitoids and Late Cretaceous granitoids in the xenoliths were entrained under the same pressure-temperature SW Japan. Data from Czamanske et al. (1981),Honma (1974), Kanisawa (1979), Murakami conditions that existed during phenocryst growth in the spessartite. (1981), Tainosho (1973, 1986), Tainosho et al. (1979), and this study. 5. K-Ar ages Leake et al., 1997), and show distinct zoning (Fig. 3C) from pleochroic brown (magnesiohornblende and magnesiohastingsite） through light K-Ar dating was carried out on hornblende separates to constrain brown and green (magnesiohornblende) to colorless (ferrohornblende). the onset of Cretaceous magmatism in the district. Fourteen new K-Ar Table 2 K-Ar age data for the Kyoto lamprophyres, hornblende peridotites, and Tamba granitoids. Sample no. Rock type Mineral Size(#) K (wt.%) Rad argon40 (10-8ccSTP/g) K-Ar age (Ma) Non rad Ar (%) Kyoto lamprophyre (stageI, Early Cretaceous) KY-501 hbl Spessartite 35°06′58" hbl 200-250 0.650 ± 0.013 263.7 ± 3.0 101.6 ± 2.3 9.4 135°36'00” KY-502 hbl Spessartite 35°06'30" hbl 200-250 0.999 ± 0.020 400.0 ± 4.7 100.4 ± 2.3 11.6 135°37'54" KY-505 cpx-hbl Spessartite 35°10'41" hbl 200-250 0.487 ± 0.024 198.5 ± 2.2 102.0 ± 5.1 11.5 135°41'54" KY-56A hbl Spessartite 35°10′33" hbl 200-250 0.402 ± 0.020 175.2 ± 2.1 109.0 ± 5.4 16.3 135°44'45" KY-56B hbl Spessartite 35°10′33" hbl 200-250 0.396 ± 0.020 168.0 ± 2.3 106.1 ± 5.3 23.9 135°44'45" U-0001 hbl Spessartite 34°52'44" hbl 200-250 0.551 ± 0.011 231.0 ± 2.7 104.9 ± 2.4 12.0 135°50'41" I-0024a hbl Spessartite 35°10′33" hbl 200-250 0.595 ± 0.010 258.0 ± 3.0 108.0 ± 5.0 74.9 135°44'45" KY-044 hbl Spessartite 34°57'06" hbl 150-200 0.695 ± 0.014 280.4 ± 3.5 101.1 ± 2.3 10.5 135°38/34" KY-044 hbl Spessartite 34°57'06" hbl 200-235 0.675 ± 0.014 265.7 ± 2.9 98.7 ± 2.2 9.6 135°38'34" Porphyry (stage Il, Late Cretaceous) KY-503 cpx-hbl Porphyry 35°07'30" hbl 200-250 0.472 ± 0.024 153.6 ± 2.8 82.0 ± 4.3 39.3 135°37'36" KY-504 cpx-hbl Porphyry 35°11'04" hbl 200-250 0.460 ± 0.023 163.5 ± 1.9 89.3 ± 4.5 13.8 135°42'24" U-0003 cpx-hbl Porphyry 34°50'27" hbl 200-250 0.508 ± 0.010 172.3 ± 1.9 85.4 ± 1.9 8.9 135°53/44" Mantlexenolith hbl Peridotite KY-56U 35°10′33" hbl 200-250 0.429 ± 0.021 195.3 ± 2.6 113.7 ± 5.7 23.4 135°44'45" KY-500 hbl Peridotite 35°10′33" hbl 100-150 0.462 ± 0.009 194.8 ± 2.2 105.5 ± 2.4 13.9 135°44'45" KY-500 hbl Peridotite 35°10′33" hbl 145-200 0.441 ± 0.009 186.6 ± 2.1 105.9 ± 2.4 13.0 135°44'45" Tamba granitoids HA-04b Granodiorite 35°10'04" hbl 200-250 0.320 ± 0.016 129.5 ± 2.4 101.5 ± 5.2 39.2 135°46′36" MO-02b Tonalite 35°08/19" bt 80-100 6.795 ± 0.136 2820 ± 29 103.9 ± 2.3 2.7 135°47'24" KTS-2b Granodiorite 35°09'05" bt 80-100 6.565 ± 0.328 2704 ± 29 103.1 ± 5.1 2.3 135°46/38" KI-002b Granodiorite 35°10'59" hbl 200-250 0.452 ± 0.023 192.6 ± 3.3 106.5 ± 5.5 36.1 135°32'59" hbl = hornblende,bt = biotite a Kimura and Kiji (1993). b Kiji et al. (1995). Table 3 114 Whole rock chemical and Sr-Nd isotope compositions of the Kyoto lamprophyres, Tamba granitoids, and peridotite xenoliths in the lamprophyre dikes. No. KY-104 KY-32 KY-009 KY-501 KY-11 KY-33 KY-34 E044 KY-1505 KY-15 KY-01b KY-04 Rock type Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. Kyoto L. H** H** H** H** I** L** L** L** L** SiO2 wt.% 46.97 47.91 51.02 51.24 46.05 47.49 48.46 48.69 50.00 50.16 51.96 52.23 TiO2 1.86 1.66 1.03 1.13 1.04 1.13 1.17 a lack of foraminifers, macrofossils and bioturbation, and a prominent positive excursion of 8l3Corg. A significant hiatus during the 1.08 1.05 0.91 1.04 Al2O3 14.05 14.22 13.23 13.58 Cenomanian mudstone-dominant unit); the Saku Formation (Middle Cenomanian-Upper Turonian sandstone-common turbidite 12.56 13.99 17.2 13.94 17.49 sandstone and mudstone, showing occasional lateral facies changes, have caused confusion regarding the lithostratigraphic 12.92 Fe203 10.08 696 8.06 8.10 12.82 11.43 11.30 8.99 9.96 9.85 9.71 8.90 Mn0 0.14 0.15 0.15 ELSEVIER 0.21 0.2 0.17 0.16 0.16 0.2 0.17 0.15 MgO 11.39 10.51 9.87 8.99 13.19 11.36 10.62 7.04 886 7.37 9.25 10.08 Cao 7.61 9.32 8.11 7.86 9.62 8.68 8.12 9.07 6.50 9.81 7.79 7.07 Na20 3.59 2.77 3.89 4.32 1.31 2.48 2.78 2.43 2.88 2.49 2.52 2.52 K20 0.94 1.82 1.69 1.49 Ryoji Wani*, Hisao Andoe * Corresponding author 1.35 0.83 2.49 1.38 1.35 2.04 LOI 2.90 3.39 2.33 2.83 4.12 3.15 3.62 4.93 2.36 0.16 3.34 2.46 P205 0.70 0.31 0.35 0.31 0.21 0.27 0.28 0.12 0.24 0.19 0.19 0.19 Total 100.23 101.75 99.73 99.98 99.50 99.82 101.86 100.56 99.44 100.15 100.12 99.60 CIPWminerals 20 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.62 0.00 m加d 1.53 2.25 1.59 2.17 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 T.Imaokaetal./Lithos 0.00 00°0 0.00 0.00 21.99 6.87 5.53 22.64 6.97 12.76 29.03 25.99 25.02 19.88 18.16 16.65 13.86 20.44 20.14 1.29 18.33 www.elsevier.com/locate/CretRes 0.00 3.12 Mg# 71.8 71.7 73.4 71.4 69.3 68.9 68.6 64.6 68.9 63.6 68.1 71.8 XRF (ppm) 217 223 160 178 311 245 245 237 203 256 219 222 Cr 378 523 571 506 1202 834 605 214 566 321 547 695 47240 43 421435 45 78 52 50 2953 88 37 51 39 184- Ni 232 14 420 219 216 236 20 119 215 187(2014)105-131 n 100 104 63 12 68 30 .5 65 Zn 105 93 77 70 110 113 102 113 86 105 84 68 Pb 2.6 7.0 9 6.1 6.4 1.2 6.2 36 2.2 8.0 5.0 1.1 ICP-MS(ppm) 181 18:3 19 19:1 156m 1517 18:2 2035m 16.5 16 164 415m 26 52 IS 69 326 685 838 263 512 619 228 802 487 252 791 Y doi:10.1016/j.cretres.2004.02.004 17.3 16.6 18.8 17.8 22.8 25 30.3 22.2 22.1 18.9 21.9 177 136 175 169 87. S6 93 97 114 72 111 105 34.5 13.8 21 19.1 7.0 8.3 9.0 4.1 9.7 6.7 13.2 8.5 S 0.1 1.7 0.3 0.6 1.0 0.9 1.4 0.9 <0.1 5.2 0.5 0.1 Ba 286 334 471 358 112 325 468 171 452 287 334 380 72.8 33.4 47.8 41.9 14.5 15.6 17 16.0 17.8 16.7 22.1 18.8 131 70.5 95.8 88.3 31.1 34.2 36 30.4 38.9 35.6 47.1 35.5 Pr 16.8 8.75 10.3 10.5 3.66 4.55 4.70 4.11 5.04 4.62 5.28 4.85 PN 64.6 33.8 39.6 41.1 15.6 19.9 diversity become apparent above the MCE horizon. In the study area, the OAE2 horizon has been well documented, and is placed in 18.1 21.5 20.7 21.1 20.8 Sm 11.1 6.31 7.23 7.44 3.7 4.82 4.69 4.08 4.61 4.87 4.53 4.65 2.95 1.75 2.01 1.90 1.09 1.31 1.35 1.22 1.24 1.36 1.08 1.33 8.08 5.11 5.25 4.42 3.8 Cretaceous Research 25 (2004) 365-390 4.79 4.58 4.37 4.60 3.95 4.32 Tb 1.05 0.69 0.66 9'0 0.62 0.65 0.80 0.73 CRETACEOUS RESEARCH 0.7 Dy 5.12 3.68 3.49 3.44 3.55 3.91 4.33 the middle part of the Saku Formation. 4.23 4.00 on the north-west Pacific margin 4.12 Ho Abstract 0.65 0.66 0.64 0.7 0.75 0.87 0.90 0.86 0.78 0.71 0.81 Er 2.33 1.73 1.83 1.72 2.02 2.1 2.53 2.63 2.46 2.31 2.09 2.35 Tm 0.312 0.229 0.256 0.249 0.303 0.315 0.352 0.380 0.371 0.330 0.311 0.35 Yb 1.83 1.49 1.61 1.59 2.00 2.09 2.25 2.50 2.35 2.06 2.02 2.26 Lu 0.256 0.236 Formation (Middle Turonian-Campanian shelf mudstone/sandstone unit), which correspond in age to the shallower facies of the 0.243 Hikagenosawa Formation consists of weakly laminated, pyrite-rich mudstone. Planktonic and calcareous benthic foraminifers are 0.329 0.355 0.411 0.341 0.330 0.318 0.353 JH 4.2 3.1 3.8 4.0 2.3 2.7 2.10 2.8 1.8 2.8 2.7 2.1 e1 2.02 6'0 1.05 1.13 0.41 0.49 0.57 0.26 0.68 0.47 0.61 0.51 Th 4.68 7.32 6.90 2.52 2.61 2.58 2.80 4.24 5.07 5.59 7.5 U 1.54 0.99 1.56 1.49 0.59 0.64 0.56 1.10 0.91 1.09 1.2 0.86 TIMS isotope ratio 87Rb/sr 0.046 0.110 units, with intercalations of six distinct stratigraphic key 0.066 0.705050 9600 0.162 0.165 87sr/6Sr 0.704588 0.705607 0.705412 0.705400 0.705260 0.704739 0.704763 ± (2o) 0.000013 100000 100000 0.000014 0.000010 0.000014 0.70510 0.000011 0.000014 Sr (105 Ma) 0.70452 0.70525 0.70529 0.70516 0.70491 0.70450 0.70452 Sr (105 Ma) 14Sm/+4Nd 2.0 10.3 12.4 13.0 11.10 7.53 1.71 2.0 10N/Nd 0.104 0.113 0.110 0.1094 0.143 0.146 is very confused. Various definitions have been proposed, 0.130 0.135 0.512657 0.512660 0.512657 0.512591 0.512477 0.512517 0.512580 0.512447 0.512549 ± (20) 0.000013 110000'0 110000'0 0.000012 0.000012 0.000014 0.000012 0.000014 0.000014 Nd, (105 Ma) 0.51259 0.51258 0.51258 0.51252 0.51238 0.51242 0.51246 0.51236 0.51246 cNd (105 Ma) 1.61 1.55 1.52 0.25 the Hakobuchi Formation (shallow-marine sandstone 1.69 0.91 2.83 160- No. KY-506A KY-14 KY-08 KY-13 KY-12 KTN-2 YA-1 HA-9 UN-1 KTS-3 TAW-1 KI-2 Rock type Kyoto L Porphyry Porphyry Porphyry Porphyry Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. L"* SiO wt.% Hokkaido is basically characterized by six alternations 53.51 53.59 55.47 56.81 53.63 57.09 61.59 63.08 63.27 63.35 63.86 1.00 1.09 1.00 0.97 0.76 1.40 1.19 0.75 0.66 0.74 0.69 AlO 12.81 15.31 15.38 14.73 15.55 15.03 15.20 15.92 15.16 16.42 15.46 15.52 T. Imooka FeOs 9.59 9.66 9.01 0.18 7.14 4.56 4.77 4.86 4.98 4.45 MnO 0.17 0.16 0.14 0.16 0.16 5.39 0.14 0.09 0.07 0.09 0.09 0.08 MgO 9.07 6.58 5.49 6.25 5.78 4.44 2.95 2.77 2.37 2.80 2.76 CaO 7.17 6.16 5.47 6.16 6.13 7.48 6.45 4.63 4.61 4.44 4.23 4.31 et al. Na;O 2.55 2.58 2.50 2.70 2.24 3.18 4.82 4.71 4.32 4.72 4.58 /Lithos KO LO1 1.93 1.63 2.12 2.54 1.60 2.27 1.45 2.17 2.35 1.89 2.04 2.66 3.13 4.30 2.33 2.67 1.76 1.94 2.62 1.23 E80 1.21 0.74 P;Os 0.14 0.17 0.24 0.16 0.13 0.51 0.44 0.25 0.28 0.30 0.28 0.22 Total 8866 8666 99.72 866 100.09 1666 99.48 966 99.51 99.98 99.75 99.25 CIPW minerals (2014） 0.55 13.22 14.8 1123 1646 2.20 .3 13.39 13.59 15.08 14.81 0.00 15.67 0.00 0.00 0.00 0.00 28.29 definitions proposed in this paper are as follows, in as- 23.59 25.87 22.53 19.39 15.78 shelf to shoreface Mikasa Formation and the overlying, 9.46 10.64 11.10 10.19 105 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 131 Mg# 68.0 61.1 59.7 62.1 61.8 59.3 58.4 59.3 57.0 52.9 56.3 58.7 XRF (ppm) V 229 225 S61 861 169 2 IS1 95 G 528 290 161 320 176 EL 108 8艺 47 2713 21 22 21010 63 68 66 15 86 23 76 6 40.6 26.7 37.7 27.8 19.3 41.4 26.2 Cu 6 10 6.1 220 27 41 n.d. 5.7 Zn 88 77 82 75 L6 52 35 45 n.d. 48 Pb 3.3 4.2 4.3 n.d. 9.9 5.1 0.7 2.8 16 0.5 ICP-MS (ppm) Ga 16 1738 17 17 167 29 21 29 21 20 24 20 Rb 41 6 69 90 49 50 > 218 69E 567 490 407 828 1070 1040 1020 1060 19.8 24.1 27.3 40.6 24.5 22.5 17.8 11.3 13.3 16.8 15.3 11.5 98 126 160 145 124 154 129 215 172 235 175 15.9 158 13.1 10 8.5 16.4 151 17.1 17.1 24.1 18.3 1.7 15.4 0.4 0.5 2.9 10.1 9.7 0.4 1.2 2.4 464 585 1030 769 659 484 697 680 600 La 21.8 20.9 32.6 25.8 19.3 40.7 36.5 51.7 40.4 67.1 46.6 40.1 (continued on next page) Table 3 (continued) No. KY-506A KY-14 KY-08 KY-13 KY-12 KTN-2 YA-1 HA-9 UN-1 KTS-3 TAW-1 KI-2 Rock type Kyoto L Porphyry Porphyry Porphyry Porphyry Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. L"* Ce 44.9 45.5 61.6 50 40.8 89 73.8 89.4 76.2 128 84.6 rimmed shelf along the Asian continental margin (Sano, Pr 9'S 5.32 6.89 4.74 11.2 8.76 668 S9'8 13.1 9.32 8.2 Nd 22.5 22.4 600 28.7 19 46.4 rounded granules to pebbles of mudstone, chert and 32.1 32.5 46.9 34 30.5 Sm 4.71 5.12 6.47 2003) (Fig. 8). A significant hiatus is present from 4.24 906 6.5 5.24 5.43 7.79 5.88 5.07 Eu the Planomalina cheniourensis toHedbergella planispira 1.38 1.75 2.06 1.22 2.34 81 pebbles-boulders of rhyolite, with subordinate, well- 1.45 1.96 1.58 1.36 Gd 4.1 4.75 5.91 7.27 4.23 6.42 4.86 4.03 99' 5.65 4.07 3.56 Tb 0.7 as much as 900 m (Fig. 6; section 10). The conglomerates 60 1.21 0.72 88'0 0.48 60 0.71 0.57 0.46 Dy 4.15 4.52 formation are usually less than 2 m thick. However, in 7.05 4.45 4.5 3.41 2.32 2.47 3.51 2.87 2.35 Ho 0.81 0.92 60 1.41 60 9'0 0.43 0.44 S9'0 0.51 0.42 Er 2.34 2.7 2.66 4.18 2.7 2.16 1.74 1.21 1.29 1.85 1.43 1.17 Tm are composed of platy glass shards, subordinate idio- 0.404 0.408 0.62 0.427 0.285 0.243 0.173 0.182 0.268 0.206 0.173 q 2.28 2.66 2.7 4.01 2.95 1.76 1.49 1.1 1.14 1.83 1.31 1.07 3 0.363 0.428 0.432 629'0 0.471 0.25 0.209 0.179 0.164 0.289 0.185 Ticinella primula planktonic foraminifera (pf) Zone to 3.2 3.6 3.5 3.0 3.5 4.8 4.1 e1 0.71 8E'0 0.51 0.61 0.57 1.11 1.34 6°0 1.2 1.64 5.2 1.33 4.2 1.38 3.7 Th 89°S 4.2 5.5 3.68 4.33 7.73 6.53 6.52 9.51 12.7 11.6 8.56 T. Ima U 1.1 60 1.4 1.44 1.5 1.79 1.79 1.28 2.50 2.75 3.32 2.49 TIMS isotope ratio 87Rb/86Sr 0.314 0.132 0.136 87s/8655 0.704595 0.703979 0.703714 ± (20) R.Takashima et al./Cretaceous Research 25(2004） 365-390 0.000014 0.000011 Lithos Sr; (105 Ma) 0.70413 0.70378 0.70351 147Sm/14Nd eSr (105 Ma) 3.56 8.45 12.3 184 373 0.113 6600 0.101 YezoGroup 0.512602 0.512724 ± (2o) 0.000014 (2014) Nd, (105 Ma) 0.51257 0.51253 εNd (105 Ma) 1.31 0.61 2.97 105131 No. UM-2 HT-2 FU-2 HA-5 MO-3 MO-20 FU-20 FU-21 KY-506B KY-02 KY-51 Rock type Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Tamba Gr. Xenolith Xenolith Xenolith Si0 wt.% 64.39 64.39 64.43 65.76 69°99 67.47 96'89 77.61 45.55 48.03 48.43 TiO2 690 0.67 0.64 0.64 0.45 0.46 0.42 0.18 0.48 0.61 blage of hard felsic volcaniclastic sandstones, tuffs and Al;O 15.79 16.47 16.03 15.41 15.57 15.52 16.27 12.45 859 7.63 10.83 Fe0 4.49 4.34 4.72 3.95 2.90 2.85 1.39 0.64 10.58 10.19 9.96 MnO 200 S00 60'0 0.07 S0'0 0.05 E00 100 0.12 0.15 0.16 MgO 2.43 1.26 1.94 2.50 1.71 1.64 0.7 0.37 25.46 22.87 14.65 deposition of the Kirigishiyama Olistostrome Member. 4.02 and macrofossils have been obtained from this forma- the lower part of the Biticinella breggiensis pf Zone, 3.94 3.47 tion, radiolarians (spumellarians） occur abundantly 3.27 68°0 4.24 5.77 8.62 NazO 4.36 3.96 4.02 4.46 4.72 4.57 4.5 2.74 duboisi (Chevalier). Agglutinated and calcareous benthic 0.5 680 KO 2.23 2.55 2.42 2.19 1.72 1.91 2.84 4.11 0.21 0.21 S90 LO1 POs 160 1.23 1.50 1.17 2.32 1.24 1.69 1.80 6.20 4.77 0.19 0.29 0.19 0.22 0.14 0.16 0.17 0.20 0.07 0.09 0.13 Total 99.57 St66 966 100.31 99.74 100.24 101.00 99.96 966 99.74 CIPV minerals Qz 17.26 19.74 19.32 18.74 22.25 23.14 23.81 44.18 000 0.00 0.00 ne 000 000 10.33 00'0 000 0.00 7.60 000 69 000 0.00 00°0 0.00 0.00 hy 10.34 7.96 9.51 2.88 1.46 assigned to the Lower-lower Upper Albian (Nishi et al. 48.58 40.66 ol 00°0 000 0.00 0.00 00°0 0.00 000 0.00 26.79 14.68 2.83 Mg# 55.4 866 48.4 59.3 57.2 57.3 54.1 57.4 (KY-2), which extends throughout Hokkaido (Table 1). thickness of the conglomeratic beds exceptionally attains CIPW minerals 64 50 5 20 1515 1250 71520m 1460 120 176 1278 06 143 112 98 5.80 17.5 29.5 1472 1235 37 1.5 1.3 19 13 35 30 19 12 LS9 20 Planktonic 57 65 Pb 5.3 5.1 2.9 1.7 4.8 2.00 3.0 4.0 1.1 n.d. 2.3 ICP-MS (ppm) 2.2 265 27 23 207 26 2.5 16 8.0 0 12.0 17.0 Y 887 924 Fm G. arca 831 848 1125 296 87.0 78.9 214 Nishiet al. (2003) 17.6 17.2 10.5 7.9 8.3 14.3 8.3 15.7 Zr 162 228 187 163 155 142 211 108 42.0 H.helvetica 73 Nb 14.2 22.6 Kanajiri 15.4 9.5 9.0 18.4 14.7 3.9 5.2 6.4 Composite 1.8 2 1.1 1.1 0.7 2.5 1.2 1.8 0.5 Ba 561 552 616 507 500 537 826 963 28 17 302 の凹 36.8 56 42.2 39.4 29.6 26.0 53.5 50.8 7.19 9.52 13.9 67.1 76.6 71.3 54.8 47.4 92.8 78.4 15.8 22.6 27.3 381 -23 8.48 7.74 5.61 5.28 10.0 2.02 2.71 3.57 27.3 39.7 30.6 27.9 20.7 19.0 Litho- 21.1 9.02 Tsukimi 14.5 Sm events 6.12 5.29 Stage 3.64 3.39 5.37 2.95 1.88 2.3 3.05 T.Imaokaet al./ Lithos1 Eu 1.33 1.6 1.35 units 1.1 60 1.45 0.85 0.494 0.625 0.799 Gd 3.59 4.14 3.82 3.15 2.65 2.28 3.34 nian 1.78 2.16 2.72 Tb 0.52 0.58 0.55 0.42 0.35 0.30 0.48 0.23 0.3 0.35 0.46 Pyrite-rich G.ferre 2.92 2.09 C. fornicata 1.57 2.62 1.30 1.75 2.1 iSst Mbr Hakkin 0.54 column 0.32 Oceanic 0.46 Santonian !Sst Mbr 0.42 Lackof 1.4 1.74 1.58 B.breggiensis 0.9 0.76 1.30 0.86 0.99 nian 1.56 Tm 0.200 0.259 0.238 0.152 0.129 0.116 0.205 0.147 0.142 0.188 0.238 Yb 1.25 1.71 1.59 0.98 0.84 0.77 1.37 1.0 0.91 1.22 1.5 184-187 (2014) 105-) 0.172 0.247 0.227 0.149 0.135 0.111 0.200 0.144 0.133 0.193 0.217 3.9 5.2 4.5 4.0 3.8 3.1 4.1 3.0 1.2 1.5 1.8 Ta 1.33 2.03 1.87 1.93 biohorizons 0.53 -4genera 1.28 0.25 0.27 0.43 Th 9.63 12.9 11.9 11.5 -22 6.44 11.7 26.5 1.75 1.84 U L. cabri 2.59 2.79 3.22 diversity 1.73 2.87 4.67 0.37 0.41 0.7 TIMS isotope ratio -131 87Rb/86Sr 0.214 0.164 0.1430 0.1791 0.704245 0.703915 0.704783 0.704773 ±(20) 0.000013 0.000012 0.000013 0.000014 Sr; (105 Ma) 0.70393 0.70367 0.70457 0.70451 εSr (105 Ma) 6.4 10.02 2.74 1.83 147Sm/144Nd 0.097 0.106 0.126 0.123 143Nd/144Nd 0.512567 0.512649 0.512406 0.512493 ±(20) 0.000016 0.000014 0.000014 0.000014 Nd; (105 Ma) 0.51250 0.51258 0.51232 0.51241 εNd (105 Ma) 0.05 1.42 3.58 1.84 H*: high-(La/Yb)n group, L** low-( La/Yb)n group. Note: CIPW norms and Mg# were calculated with 15 at.% of totaliron as Fe3+ and 85% as Fe2+, as recommended by Verma et al. (2002) and Verma (2003). 三 118 T. Imaoka et al./Lithos 184-187(2014)105-131 20 1500 Al2O3 (wt.%) Cr (ppm) oTamba Lamprophyre 18 A 1200 + Tamba granitoids 16 900 porphyry (LateKrt) A 9 14 88 600- 。 12 300 10 0- 6 500 Na20 (wt.%) Ni (ppm) 400 300- 。 200- 2 。 100- 80 0+ 2000 K20 (wt.%) Ba (ppm) 4- + High-K 1500 3 1000- # + 2- Q ++++ +4 +# Medium-K Low-K 500 o 15 150 0 MgO (wt.%) Rb (ppm) △ 10- 100- # # 50 5- +# LSA HSA 1.5 1200 MgO/Fe2O3 Sr (ppm) 1000 800 600 @ 0.5 400 品 O4 200 88 0 -0 40 P205 (wt.%) Nb (ppm) o HNB 20 NEB 8 8 10 LSA o HSA 0+ 0+ 45 50 55 60 65 70 75 80 45 50 55 60 65 70 75 80 SiO2 (wt.%) SiO2 (wt.%) Fig.7.Harkervariation diagramsshowing themajorand tracelement variationsintheKyotolamprophyres,porphyries,and theTamba granitoids.DatafortheTambagranitoids from Kiji et al. (2000) are also plotted. K20-SiO2 classification diagram with IUGS fields after Le Maitre (2002). HSA = high-silica adakite,LSA = low-silica adakite (Martin et al.,2005), HNB = - high-Nb basalt (Defant et al., 1992), and NEB = Nb-enriched basalt (Defant et al., 1991). ages and four previously published ages of Tamba granitoids (Kiji et al., 82.0 ± 4.3 Ma). Hornblendes from the hornblende peridotite xenolith 1995) are listed in Table 2. Dike rocks can be clearly divided into two yielded ages of 105.5 ± 2.4 Ma and 105.9 ± 2.4 Ma, identical with groups: the Early Cretaceous Kyoto lamprophyres (109.0 ± 5.4 Ma to ( 1 1 o) 5 n 98.7 ± 2.2 Ma) and Late Cretaceous porphyry dikes (89.3 ± 4.5 Ma to analytical error, suggesting a close relationship between the T.Imaoka et al./Lithos 184-187(2014)105-131 119 spessartite and the hornblende peridotite xenolith. The Tamba granit- low-(La/Yb)n group (3.6-6.1) with less steep REE patterns (Fig. 9). The oids gave hornblende K-Ar ages of 106.5 ± 5.5 and 101.5 ± 5.2 Ma, high-(La/Yb)n group shows steeper REE patterns than average ocean is- and biotite ages of 103.9 ± 2.3 and 103.1 ± 5.1 Ma (Table 2; Kiji land basalt (OIB) (Sun and McDonough, 1989), and compared with the et al., 1995). low-(La/Yb)n group it has lower Yb concentrations (1.6-1.8 ppm com- From these data, it is evident that the Early Cretaceous Kyoto pared with 2.0-2.5 ppm). CIPW norms indicate that the high-(La/Yb)n lamprophyres, the hornblende peridotite xenoliths, and the Tamba group is slightly nepheline normative (ne = 1.5-2.3%), whereas the granitoids overlap almost completely in their ages, and yield a tight low-(La/Yb)n group is olivine normative (ol = 1.3-20.4) to SiO2- cluster at 109-99 Ma, hence confirming the existence of Early Creta- n (-6i = y 'io~ = ) z ceous adakitic and lamprophyric magmatism in the Kinki district of porphyries are Qz normative subalkaline rocks (Qz = 22.5-27.6). SW Japan. Primitive-mantle-normalized diagrams (Fig. 9; Sun and McDonough, 1989) show that the absolute abundances of trace elements are highly 6. Geochemistry variable, although both groups have similar trace element characteristics. The analyzed samples are characterized by enrichments of Ba, Th, U, and 6.1. Major and trace element chemistry Pb relative to average OIB (Sun and McDonough, 1989), and slight deple- tion in Nb (despite overall high Nb concentrations), Ta, Ti, HREEs, and Y, 6.1.1. Kyoto lamprophyres (Early Cretaceous） and porphyries (Late and thus exhibit the established characteristics of subduction-related Cretaceous) magmas. Representative analyses are listed in Table 3. LOI shows a wide varia- The lamprophyres are also characterized by elevated levels of high tion in volatile content from 2.3 to 4.1 wt.%. The highest LOI value reflects field strength elements (HFSEs) such as TiO2, Nb, and Zr (up to the high percentage of carbonate in the groundmass, but we note that 1.68 wt.%, 34.5 ppm, and 250 ppm, respectively). Niobium contents of even higher volatile contents (up to 5-6 wt.%) are a common feature of the high- and low-(La/Yb)n groups are 13.8-34.5 (avg. 22.1) ppm and lamprophyres (Rock, 1991). Because of these high LOI values, the major 4.1-13.2 (avg. 8.3) ppm, respectively; these averages roughly corre- element contents have been recalculated on an anhydrous basis. The spond to the values in high-Nb basalt (HNB: Nb > 20 ppm; Defant Kyoto lamprophyres display the very specific geochemical characteristics et al., 1992) and Nb-enriched basalt (NEB: Nb = 6-20 ppm; Defant of primitive high-Mg basaltic to basaltic andesite compositions, with et al., 1991), respectively (Fig. 7). The high Nb/Y ratios (0.8-1.5) confirm lower SiO2 (anhydrous basis) contents of 48-54 wt.%, higher MgO con- the alkaline character of these ne-normative rocks, and together with the tents (5.4-14.1 wt.%), and higher Mg0/Fe203* ratios (1.0-1.2) than in Zr/TiO2 ratios they make the rocks similar to alkali basalt (Winchester the porphyries (Fig. 7). High concentrations of compatible trace elements and Floyd, 1977). (e.g, Cr = 214-1202 ppm, Ni = 55-420 ppm, Co = 37-78 ppm, and The lamprophyres of both groups also have high Th/Ta ratios V = 160-311 ppm) sugest that the lamprophyres did not experience (3.9-14.8), and thus plot in the volcanic-arc basalt field of the Th-Hf/3-Ta diagram (Fig. 10A) of Wood (1980). The lamprophyres are they therefore most probably represent the mantle-derived primitive also characterized by low Nb/U ratios (12.8-22.4; Table 3), significantly magma. lower than those of MORB and OIB (~47) (Hofmann et al., 1986). These In terms of the total alkali-silica (TAS) diagram of Le Maitre (2002) chemical characteristics are consistent with a subduction-related envi- (Fig. 8), the lamprophyres have a wide range of NazO + K20 contents ronment of magma generation (e.g., Pearce, 1983; Pearce and Peate, (1.1-6.1 wt.%), and plot in both the basaltic trachyandesite field of the 1995). alkaline-rock series, and the basalt-basaltic andesite field of the The high- and low-(La/Yb)n groups have also been plotted on the subalkaline-rock series. On the other hand, the analyzed porphyries all Th/Yb vs. Nb/Yb diagram (Fig. 10B) of Pearce (1983, 2008). These ratios plot in the subalkaline-rock series (after Irvine and Baragar, 1971). are almost independent of fractional crystallization and/or partial melt- The lamprophyres are sodic, as NazO always exceeds K2O, and K20/ ing with pyroxenes and feldspars as the dominant phases, and thus NazO ratios are 0.23-0.97. highlight source variations and crustal assimilation. Basaltic magmas The REE patterns of the Kyoto lamprophyres are characterized by derived from the mantle lithosphere lie within or close to a diagonal enrichment in light REEs (LREEs), and are similar to those reported for MORB-OIB array defined by constant Th/Nb ratios. Source region meta- lamprophyres by Rock (1991). They have large variations in LREEs in re- somatism as a result of subduction processes, however, produces an en- lation to the heavy REEs (HREEs), and can be separated into two groups: richment of Th with respect to Nb, and hence Th/Yb ratios greater than a high-(La/Yb)n group (12.5-22.1) with very steep REE patterns, and a Nb/Yb, because subduction components generally carry Th but not Nb or Yb. Crustal contamination may also increase Th/Yb ratios relative to Ta/Yb ratios due to higher abundances of Th relative to Nb in crustal rocks (Pearce, 1983). Fig. 10B shows that both the high- and low- 10 (La/Yb)n groups are clearly displaced above the MORB-OIB array. Al- basaltictrachyandesite trachytel h trachy- trachydacite though the effects of crustal contamination on magma compositions are rachy difficult to distinguish from those of sediment involvement (metasoma- rhyolite tism) as a result of subduction processes, the significantly high Th/Yb ratios for the most basic samples of the high- and low-(La/Yb)n groups, dacite and the low Si02 wt.%. of 48.26 and 48.28, respectively, are more likely andesite - psn q a s a a o C oKyoto Lamprophyres tamination. Plank (2005) showed that Th/La ratios are low (<0.2） in saltic + Tamba granitoids Oceanic basalts (OIB and MORB), but vary in arc basalts due to mixing basalt desite porphyries (Late-K) between mantle (low Th/La) and subducted sediments (high Th/La). Both the high- and low-(La/Yb)n groups have higher Th/La ratios of 45 50 55 60 65 70 75 80 0.11-0.40; this indicates that both the magmas inherited components SiO2 (wt.%) from subducted sediments. High ratios of both Nb/Yb and Th/Yb relative to N-MORB suggest that the magma was generated either by melting of Fig. 8. TAS (total alkalis vs. SiO2) diagram for the Kyoto lamprophyres, Tamba granitoids, and Late Cretaceous porphyries. The framework is after Le Maitre (2002). The broken bearing mantle source, or by a combination of these processes, as will lines a (after Kuno, 1966) and b (after Irvine and Baragar, 1971) show the boundary be- tween alkaline and subalkaline (tholeitic) rocks. be described later. 120 T. Imaoka et al./Lithos 184-187(2014)105-131 In summary, the high- and low-(La/Yb)n groups have primitive 2000; Tsuchiya et al., 2005; Yogodzinski et al., 1995). This is also re- high-Mg basalt and basaltic andesite compositions, suggesting insignif- sponsible for variable SiO2 contents (Martin et al., 2005). They have icant fractionation or crustal assimilation in their petrogenesis. They high Al203 contents of 15.0-16.5 wt.%, averaging 15.5 wt.%. The also carry a clear subduction-related geochemical signature. Neverthe- Tamba granitoids are characterized by high Sr (827-1125 ppm) con- less, the two groups have some distinctive geochemical characteristics: tents but low Y (mostly <18 ppm) and Yb (mostly <1.8 ppm). They the high-(La/Yb)n group has higher Nb/Yb and Th/Yb ratios relative to exhibit high Sr/Y (41-105) and (La/Yb)n (23-47) ratios and steep the low-(La/Yb)n group. These data suggest that the two groups repre- REE patterns lacking Eu anomalies (Figs. 9, 11). They are strongly sent partial melts from different mantle sources. sodic, and Na2O always exceeds K2O, with K2O/Na2O ratios of 0.30-0.70, except for the most silicic sample which has a ratio of 6.1.2. Tamba granitoids 1.5. On the SiO2 vs. K20 diagram (Fig. 7), many of the samples plot Analyses of the Tamba granitoids are given in Table 3. The rocks within the medium-K field (Le Maitre, 2002). have tonalitic to granitic compositions, with SiO2 contents (anhy- These geochemical signatures are adakitic, as pointed out by Kiji drous basis) ranging from 54 to 78 wt.%, Mg0 from 0.4 to 5.9 wt.%, et al. (2000), and the rocks plot in the high silica adakite (HSA) field and high Cr (up to 108 ppm), Ni (up to 58 ppm) and Co (up to (Figs. 7, 11) of Martin et al. (2005), except for the high Nb values 143 ppm) contents. These elevated concentrations of Cr, Ni, and Co, (Fig. 7). The primitive-mantle-normalized patterns contain negative as well as MgO (Fig. 6) suggest the interaction of the ascending Nb, Ta, Ce, P, and Ti anomalies (Fig. 9), similar to those of Early Creta- magma as it rose through a zone of mantle peridotite en route to its ceous adakitic granites (central facies） from the Kitakami Mountains emplacement in the crust (e.g., Kelemen et al., 2003; Kiji et al., of NE Japan (Tsuchiya et al., 2007). 1000 1000 Kyoto lamprophyres (High-(La/Yb)n group) Kyoto lamprophyres (High-(La/Yb)n group) zed ne normative : spessartite OIB 100 100 10 10 nitive ne normative : spessartite Prim OIB La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Rb Ba Th U Nb Ta La Ce Pb Sr Nd P Sm Zr Hf Eu Gd Ti Tb Y Yb Lu 1000 10000 Kyoto lamprophyres (Low-(La/Yb)n group) Kyoto lamprophyres (Low-(La/Yb)n group) & peridotite xenoliths & peridotite xenoliths 1000 hy normative: spessartite hy normative:spessartite 100 - - peridotite xenolith -peridotite xenolith 100 10 10 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Rb Ba Th U Nb Ta La Ce Pb Sr Nd P Sm Zr Hf Eu Gd Ti Tb Y Yb Lu 1000 10000 porphyries porphyries zed 1000 100 100 10 10 C La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu RbBaThUNbTaLaCePbSrNdPSmZrHfEuGdTiTbYYbLu 1000 10000 Tamba granitoids Tambagranitoids 100 10 10 C 0.1 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Rb Ba Th U Nb Ta La Ce Pb Sr Nd P Sm Zr Hf Eu Gd Ti Tb Y Yb Lu Fig. 9. Chondrite-normalized REE patterns and primitive-mantle-normalized multi-element diagrams for the Kyoto lamprophyres, peridotite xenoliths, porphyries, and Tamba granitoids. The chondrite values are from Anders and Grevesse (1989). The primitive mantle values and the average compositions of ocean-island basalt (OIB) are from Sun and McDonough (1989). T.Imaoka et al./Lithos 184-187(2014)105-131 ments in certain outcrops. In the Yezo Group, the Hf/3 B redefined as the Soashibetsugawa Formation of the dno1a u(q人/eT)-uH Th/Yb O Low-(La/Yb)n group enrichr N-MORB E-MORB -MORB VAB &WPT 0.1 N-MORB ments (origin of the TOC) were transported from the Array of basalts from non-subduction setting 0.01 Th tion (Fig. 8). This excursion interval could be correlatable abrupt, prolonged positive one (>2%o), are recorded in 1 10 100 Livello Selli in Italy and Fischschiefer in Germany, is Fig. 10. (A) Discrimination diagrams (Hf/3-Th-Ta) for the Kyoto lamprophyres after Wood (1980). WPA = within-plate alkali basalts, WPT = within-plate tholeites, and VAB = volcanic- arc basalts. (B) Nb/Yb vs. Th/Yb plots (Pearce, 2008) for the Kyoto lamprophyres. 6.1.3.Hornblendeperidotite xenoliths peridotite xenolith are 101 to 103 times greater than in other sub-arc pri- The hornblende peridotite xenolith in the spessartite has higher con- mary and metasomatized peridotite xenoliths, such as those in the tents of SiO2 (46.2-50.0 wt.%), Ca0 (4.5-9.5 wt.%),Al203 (7.1-9.2 wt.%), Avacha Volcano, Southern Kamchatka (Ishimaru et al., 2007), and in and Fe20* (10.5-11.6 wt.%), and lower contents of MgO (29.7, Ichinomegata and Kurose in Japan (Abe and Yamamoto, 1999). Some 26.9 wt.%) than typical abyssal peridotite (Niu et al., 1997), and these large ion lithophile elements (LILEs) such as Ba (26-172 ppm), Rb characteristics point to the highly fertile nature of the peridotite (2-6.7 ppm), Sr (89-120 ppm), U (0.4 ppm), and Th (1.8 ppm), as well (Table 3). It should be noted that metasomatic minerals such as Ca- as the REEs, occur in very high abundances in our peridotite samples. amphibole make up as much as 35 vol.% of the analyzed samples, and The hornblende peridotite xenolith exhibits REE and multi-element pat- their presence significantly alters the major-element composition of the terns that are similar to those of the host spessartite, and are somewhat primary bulk-rock. Abundances of some trace elements in the hornblende LREE-enriched with (La/Yb)n ratios of around 4.4 (Fig. 9). 120 lower Albian Niveau Kilian, Paquier and Leenhardt in JK (ppm) Sr/Y 0 Kyoto lamprophyres LSA 100- Tamba granitoids HSA porphyries (Late-K) LSA 80- westward Asian continental margin. Therefore, as the 60- 40 10000 the Vocontian Basin, and Livellos Monte Nerone and 20- △ fo ammonite Zone, in Tethyan-Atlantic regions (Gale 0 10 20 30 40 50 0 interval in the study area is very thick, about 110 m 150 Y (ppm) Rb (ppm) 1500 Sr (ppm) 10- Cr/Ni LSA 8 curred within the basal part of the Leupoldina cabri pf 6 HSA 4 500 HSA # LSA Q 5 10 15 0 0.5 1 1.5 CaO + Na2O (wt.%) TiO2 (wt.%) Fig. 11. Sr/Y vs. Y, K vs. Rb, Sr vs. CaO + Na2O, and TiO2 vs. Cr/Ni plots comparing the adakitic Tamba granitoids with HSA and LSA (Martin et al., 2005). 122 T.Imaoka et al./Lithos 184-187(2014)105-131 ●High-(La/Yb)n group primitive nature of the lamprophyres, together with their enriched 0 Low-(La/Yb)n group LILEs and isotopic signatures, suggest source-related features rather + Tamba granitoids than crystal fractionation; in other words, a relatively unmodified pri- ·peridotite xenoliths mary melt (small-volume partial melt) derived from an enriched man- 4- tle beneath the Kinki district. 2- 7.1.1.2. Mantle background. Sub-arc mantle metasomatism by slab- eNd 0- derived melts produces diagnostic HFSE-rich metasomatic minerals -2- such as pargasite, phlogopite, and ilmenite (Kepezhinskas et al., 1995, 1996; Prouteau et al, 2001; Rapp et al., 1999; Sajona et al., 1993, -4- 1996; Sen and Dunn, 1994; Zanetti et al, 1999). The magma of the -9- Kyoto lamprophyre was rich in water, alkalis, and perhaps other vola- tiles, as indicated by the abundance of hydrous minerals such as the -8- hornblende in the spessartites. In the lamprophyres, clinopyroxene -10- e n a s 0.7020.7030.7040.7050.7060.7070.7080.709 Na metasomatism (Kepezhinskas et al., 1995). Ti, Na, K, and HFSE-rich SrI amphiboles (Fig. 4; Table 1A; Appendix 1) might also be the products of such mantle metasomatism. High Th/Yb (Fig. 10B) and Th/La ratios hornblende peridotite xenoliths in the Kinki district, SWJapan. The field of Late Cretaceous of the lamprophyres also indicate subduction-related enrichment of granitoids in the same district (Tainosho et al, 1999) is also shown for comparison. The solid line is the mixing trend of an initial magma (Tamba granitoid KI-2: above, it is thought that the mantle beneath the Kinki district has been Sr = 1060 ppm; Nd = 30.5 ppm; Srl (105 Ma) = 0.70351; εNd (105 Ma) = 2.97,this metasomatized to various degrees by slab melts. Other direct evidence Sr = 110 ppm; Nd = 24.3 ppm; Srl (105 Ma) = 0.72648; &Nd (105 Ma) = - 15.8; of mantle metasomatism is given by the hornblende peridotite xenolith, Kagami et al, 2006; Kawano et al, 2006). which is a fragment of the lithospheric mantle beneath the forearc of Early Cretaceous SW Japan. Microscopic and geochemical features of the xenolith indicate that its hornblendes represent a reaction between 6.2. Sr-Nd isotopes a dunitic cumulate from a Mg- and Cr-rich melt and an evolved hydrous high-Mg melt (Ishimaru et al., 2009). The weight of evidence, therefore, The Sr-Nd isotope data are listed in Table 3. All our isotope data for confirms the presence of a mantle metasomatized by subduction- the Kyoto lamprophyres and the Tamba granitoids, and the peridotite related material. xenoliths were age- corrected to 105 Ma, based on our K-Ar age data Ishimaru et al. (2009) pointed out that the chondrite-normalized (Table 2). The Nd isotope system is resistant to alteration processes REE patterns and the N-MORB normalized multi-elements patterns of (e.g, Faure and Mensing, 2005), and the isotope ratios in the Kyoto clinopyroxene and hornblende in the peridotite xenolith are quite sim- lamprophyres should thus represent primary compositions. The high- ilar to those in our low-(La/Yb)n group host spessartites, and they con- (La/Yb)n group is characterized by positive eNd(105 Ma) values of 0.25 cluded that the hornblende peridotite is cognate with the host and 1.61, whereas the low-(La/Yb)n group has lower eNd(105 Ma) values spessartite. Our analyses show that the amphibole chemistry and the n o sn ( ) n ( ) - p - Sr-Nd isotope ratios in the mantle peridotite xenoliths overlap locally between the groups. The Tamba granitoids are characterized by low with the low-(La/Yb)n group lamprophyres, thus reconfirming the cog- e ‘o 0 5) ( 1） 1 J m p s nate relationship between xenolith and host lamprophyre. Nd(105 Ma) values that range from -0.1 to + 3.0. These values differ Using our K-Ar dating results, we can constrain the age relationships from those in the Kyoto lamprophyres. The hornblende peridotite xeno- between melting of the metasomatized mantle and the metasomatic lith samples have relatively constant Sr; (105 Ma) values of 0.7045 and events during subduction. In our case, the adakitic magmatism of the 0.7046, similar to CHUR, and εNd(105 Ma) values of -1.8 and -3.6. Tamba granitoids was contemporaneous with the intrusion of Kyoto All the ca. 105 Ma igneous rocks show depleted values relative to Late lamprophyres, all of which occurred within a relatively short time peri- Cretaceous granitoids in the same district (Tainosho et al., 1999). od from 109 to 99 Ma. The xenolith hornblende age of 106 Ma con- strains the timing of the hornblende-forming reaction between olivine 7. Discussion and hydrous basaltic-andesitic melts, described by Ishimaru et al. (2009), because the age is within the 109-99 Ma range of ages for the 7.1. Petrogenesis of the Kyoto lamprophyres Kyoto lamprophyres and the Tamba granitoids. HFSEs such as Zr and Hf can also provide clues about the mantle region 7.1.1. Constraints on the Kyoto lamprophyres source of the lamprophyres. Zr/Hf ratios are higher in the high-(La/Yb)n group (42.1-46.1, avg. = 43.6) than in the low-(La/Yb)n group (34.8- 7.1.1.1. Source composition. The high Mg# values (72-83), high Cr 46.2, avg. = 39.6). The ratios correlate positively with the degree of silica (161-1240 ppm), Ni (up to 395 ppm), and Co (40-117 ppm) contents, undersaturation in the basaltic rocks, and reflect compositional heteroge- and the relatively low SiO2 contents (<54 wt.%) that characterize all the neities in the source mantle that were caused by metasomatism (Dupuy Kyoto lamprophyres most probably represent mantle-derived primitive et al., 1992), and this is because there are slab controls on Zr and Hf in magma produced by partial melting of peridotitic mantle. These values arc magmas (Gomez-Tuena et al., 2011). The differences in the Zr/Hf ra- can be in equilibrium with mantle olivine with a composition of Fo90 tios may reflect differences in the mantle sources from which the high- (assuming the Fe2+ /total Fe atomic ratio is 0.85), which is equivalent and low-(La/Yb)n groups were derived. to the Fo content (94-87) of the olivine in the peridotite xenolith. Pos- Experimental studies (e.g, Kushiro, 2001; Wilson, 1989; Yoder, itive correlations between MgO and Ni (r = 0.73) and MgO and Cr 1976) led to the idea that alkali basalts can be generated by low degrees of partial melting of a garnet lherzolite source. The low HREE and Y con- tent with olivine + chromite fractionation from the parental melts, or tents and the wide range of LREE/HREE ratio in basaltic liquids should with differing degrees of partial melting of the source rocks. The reflect the presence of garnet in the source region. La/Yb ratios in partic- lamprophyres are also characterized by high LREE, LILE, and HFSE con- tents, together with negative Nb-Ta anomalies (Fig. 9). Thus, the rather lherzolite sources, because Yb is compatible with garnet but not with T.Imaoka et al./Lithos 184-187(2014)105-131 123 clinopyroxene. As described above, the Kyoto lamprophyres can be di- 100： La/Yb vided into two groups based on (La/Yb)n ratios, a high-(La/Yb)n ratio Garnet peridotite Q.1 group and a low-(La/Yb)n ratio group. The former group has higher (La/Yb)n ratios (12.5-22.1), and lower Yb (1.5-1.8 ppm) contents than 0.1 the low-(La/Yb)n group which has (La/Yb)n ratios of 3.6-6.1, and Yb con- tents of 2.0-2.4 ppm, all of which suggests different source regions. 10 Moreover, the former has higher LILE and HFSE contents. Lanthanum is a highly incompatible element, but abundances are not significantly af- Garnet-amphibole peridotite fected by the source mineralogy (e.g. garnet or spinel); La contents 0.1 thus provide information on the bulk chemical composition of the Spinel amphibole peridotite source. 10 20 30 40 50 60 7.1.2. Trace-element modeling of petrogenesis La (ppm) To evaluate these ideas we modeled LILE, REE and HFSE abundances and ratios to constrain the characteristics of the source of the Kyoto 12 Yb (ppm) lamprophyres in terms of source mineralogy, LILE, REE and HFSE con- dno.1s u(q人/eT)-u3!H centrations, and degree of partial melting. Modeling was performed 10 0 Low-(La/Yb)n group tite using the non-modal batch melting equations of Shaw (1970). Enrich- ment of LILEs, LREEs and HFSEs in the Kyoto lamprophyres, together with Zr/Nb and Yb/Ta ratios that are lower than those of normal MORB, suggests the involvement of an enriched mantle source compo- nent. Therefore, we use reference compositions of E-MORB calculat- ed for the Mid Okinawa Trough in the same western Pacific region Garnet-amphibole peridotite (Shinjo et al., 1999; Wang et al., 2004). Our work suggests that the low-(La/Yb)n group lamprophyres originated from a metasomatized . ? mantle source, and direct evidence from the mantle xenoliths points Garnet peridotite 0.1 to amphibole in the source region. Hence, four melting curves in- volving garnet-amphibole peridotite, garnet peridotite, spinel-am- 5 10 15 20 25 30 35 40 phibole peridotite and spinel peridotite are examined. Nb (ppm) In Fig. 13A, the garnet-dependent ratio La/Yb was plotted against La. The Kyoto lamprophyres that belong to the high (La/Yb)n group lie on Fig. 13. (A) La vs. La/Yb, and (B) Nb vs. Yb diagrams illustrating geochemical modeling the melting curve for garnet peridotite at low degrees (<2%) of melting, results and the geochemical compositions of the Kyoto lamprophyres.Non-modal batch melting curves are presented for garnet peridotite (Olivineo.6- and plot far from the spinel peridotite and spinel-amphibole peridotite Orthopyroxeneo.2Clinopyroxeneo.1Garneto.1: Johnson, 1998), garnet amphibole peri- melting curves (Fig. 13A). These features suggest the high La/Yb ratio dotite (Olo.55Opxo.22Cpxo.15Grto.osAmpo.o1: Barry et al., 2003） and spinel peridotite and enrichment of incompatible elements may largely have resulted (Ol0.53Opxo.27Cpxo.17Spo.03: Johnson, 1998). Smallticks with numbers represent the from low degrees of partial melting of garnet peridotite. In contrast, degree of partial melting (%). Normative weight fractions of mineral i in the partial melts are: Olo.1Opxo.18Cpxo.3Grto.42 (Johnson, 1998), Olo.05Opxo.osCpxo.3Grto.2 Ampo.1 the low-(La/Yb)n group lamprophyres are displaced from the garnet pe- (Barry et al., 2003), and Ol-0.40pxo.3Cpxo.9Spo.2 (Kelemen et al., 1993). Mineral-melt ridotite melting curve, and plots between the garnet peridotite and gar- distribution coeffcients used for the modelings are from McKenzie and O'Nions (1991 net-amphibole peridotite melting curves in their higher-degree of 1995) for REE, and Green (1994) for Nb.Source compositions: enriched mantle calculated melting regions of about 2-10%. from basalt (La 0.74 ppm, Yb 0.67 ppm, Nb 0.5 ppm; Wang et al., 2004). As described above, a Nb-rich component was involved in the petro- genesis of the lamprophyre. Petrogenesis of Nb-rich basalts remains controversial, as summarized by Hastie et al. (2011). The Nb- Taking into consideration the positive ENd(T) values (Fig. 12) and on sade dnor a dnra "(ae)-t a go au pa be the consequence of a contribution from an ocean island basalt n n s (OIB)-type mantle source (e.g, Castillo, 2008; Gazel et al, 2011; metasomatized mantle wedge with a garnet residue, and the group Reagan and Gill, 1989) or a mantle wedge that was metasomatized by therefore originated from within the garnet peridotite stability field at adakites (blends of slab and mantle materials) (Bourdon et al., 2002; depths greater than 70-80 km. In contrast, the low-(La/Yb)n group is Defant and Drummond, 1993; Defant et al., 1992; Kepezhinskas et al., displaced from the melting curve of garnet peridotite and plots between 1996; Sajona et al., 1993, 1994; Smithies et al., 2005; Straub et al., the garnet peridotite and garnet-amphibole peridotite melting curves ssod s n e na in their higher-degree of melting regions of about up to 10%. This rules the results of modeling in terms of concentrations of the HFSE, Nb and out the possibility that the low-(La/Yb)n group represents a different the garnet-dependent Yb concentration (Fig. 13B). The high-(La/Yb)n degree of partial melting of the same sources as the high-(La/Yb)n o l suu q mo e qn yu ym saudoi dnoa group. This is also indicated by the differing Sr-Nd isotope relations in the garnet peridotite melting curve between 0.1 and 3%, whereas the the two groups of lamprophyres (Fig. 12), because different degrees of low-(La/Yb)n group lamprophyre is again displaced to higher Yb, and melting do not cause variations in radiogenic isotope compositions, plots between the garnet peridotite and garnet-amphibole peridotite and the nature of the source exerts a fundamental control on the initial melting curves in their higher-degree (2-10%) melting regions. In Sr-Nd isotope ratios. As a consequence, the low-(La/Yb)n group could both diagrams (Fig. 13A, B) neither the spinel peridotite and spinel am- have been generated by mixing of a 2%-10% garnet-amphibole perido- phibole peridotite melting curves can explain the LREE/HREE ratios and tite melt with a 2%-5% garnet peridotite melt. This scenario is consistent LILE and HFSE concentrations of the high- and low-(La/Yb)n group with an amphibole-bearing mantle, as indicated by the presence of the lamprophyres. Therefore, the Nb-rich high-(La/Yb)n group could be hornblende peridotite xenolith. Both melts require residual garnet in generated by a low degree of partial melting of a garnet peridotite of a the source region. A melting column with the lower end located in the E-MORB composition without the involvement of an OIB-type source garnet-stability field and the upper end situated in the garnet amphi- in the genesis of the high-(La/Yb)n group. bole peridotite field would be a plausible explanation. 124 T. Imaoka et al./Lithos 184-187(2014)105-131 Taking into account the silica-saturation characteristics of the low- vary from 0.7035 to 0.7039 and - 0.1 to + 3.0, respectively, which are (La/Yb)n group of rocks (hy-normative), it is likely that this group repre- higher in the case of Sr and lower in the case of eNd(105 Ma) compared sents a slightly higher degree of melting, with the melts equilibrating e with mantle at shallower depths than the slightly undersaturated slab sediments were incorporated into the adakites during slab melting. high-(La/Yb)n group (ne-normative). These inferences are consistent It has already been noted by Kiji et al. (2000) that the granitoids often contain small subangular xenoliths of basement sedimentary pressures between 1 and 3 GPa (Hirose and Kushiro, 1993; Kushiro, rocks up to 7 cm in size. Tentative calculations for mixing the most de- 1994, 2001; Takahashi and Kushiro, 1983; Walter, 1998), which cover pleted Tamba granitoids with basement mudstone indicate that a max- the petrogenetic conditions and lamprophyric compositions of our sam- imum of 10%-15% mudstone mixed with the granitic magma creates the ples. To summarize, the silica contents of basaltic magmas decrease observed isotope ratios (Fig. 12). However, the estimated amounts of with increasing pressure, alkali contents decrease with increasing de- mudstone are too large relative to field occurrences. Furthermore, the grees of partial melting as low-(La/Yb)n group, and a small degree of involvement of mudstone should also cause enrichment of LILEs such partial melting at high pressures produces alkaline magmas with nor- as Pb, which is not detected even in the most isotopically enriched sam- mative nepheline as high-(La/Yb)n group. ple HA-5. These features preclude crustal contamination as a major pro- As noted above, the source was metasomatized, and so would have cess for their trace elements and isotopic variations. been accompanied by fluid and melt interactions with the mantle. En- High concentrations of HFSEs are another characteristic of the richment in fluid-mobile elements such as the LILE, and melt-mobile el- Tamba granitoids. The relationships between the contents of HFSEs ements such as Th, Nb, Ta, Zr, Hf, and LREEs, might be expected as a such as TiO2, Nb, and Ta, and the melt compositional parameter, FM result of small-scale mantle heterogeneities. These processes may be (melt basicity as expressed by increased FM), where FM is given as: reflected in the small differences in Sr-Nd isotope ratios and the LREE 1/Si·[Na + K + 2(Ca + Fe + Mg)]/Al (Ryerson and Watson, 1987), and LILE compositions of the high- and low-(La/Yb)n groups. These are shown in Fig. 14. The Tamba granitoids have high Nb and Ta con- dp e pnd aq r n psa a sn tents (up to 24 ppm and 2.0 ppm, respectively) relative to those in mantle wedge, and small-scale vertical heterogeneities in the mantle more typical high-silica adakites (Fig. 7) and adakites drawn from the wedge due to the metasomatism are therefore to be expected beneath the Kinki district of SW Japan. 2.0- Considering the above, we propose that the high-(La/Yb)n group of TiO2 lamprophyres originated from small degrees of partial melting of an enriched metasomatized mantle wedge within the garnet stability 1.5 field at depths of ≥70 km, and that the low-(La/Yb)n group of +$ lamprophyres originated from a different mantle source by a relatively 1.0 large degree of partial melting at relatively shallow parts of mantle wedge. We thus conclude that the magmatic diversity of the Kyoto +++ lamprophyres derives primarily from a heterogeneous mantle source 0.5 that has been variably affected by the results of subduction. 7.2. Petrogenesis of the Tamba granitoids (adakites) 0.0+ 1 2 3 4 5 The geochemical characteristics of the Tamba granitoids are very 25 FM Nb similar to those of typical adakites thought to have been produced by partial melting of subducted oceanic crust (Defant and Drummond, 20- 1990). The Tamba granitoids are enriched in Sr and lack or have negli- gible Eu anomalies (Figs. 7, 9) indicating that the source residue was 15 plagioclase-free. Depletion of HREEs and Y in the adakites requires melting of mafic source rocks within the garnet stability field under 10 eclogite-facies conditions (Defant and Drummond, 1990). In terms of their Sr-Nd isotope compositions, adakitic rocks can be produced from such a garnet-bearing source by partial melting of the subducted ocean slab (Defant and Drummond, 1990). The HREE patterns of the odu aru e sd ne e sisns s pe das ae sae 1 2 3 4 S role than amphibole during partial melting, and that eclogite rather 4.0- FM than garnet-amphibolite is likely to be the source lithology. If amphi- Ta bole is residual in the source, it will induce concave-upward patterns between the middle and heavy REEs (Rollinson, 1993). 3.0 + Tamba granitoids Defant et al. (2002) noted that the adakites display the following MORB-like Sr-Nd isotope characteristics. Elevated concentrations of 2.0 LILEs such as Rb, Ba, Th, and U, and their wide variations (Fig. 9) may be caused by the incorporation of sediment-derived material into the 1.0 adakitic melts during slab melting. Negative Ce anomalies, quantified by using the N-MORB normalized Ce/Ce* ratio of Ce/(La + Pr)/2), ..· have been reported from island arc lavas, and Ce anomalies in conver- 0.0 gent margin magmas have been shown to indicate contributions from 2 3 5 subducted sedimentary material inherited from seawater (e.g., Class FM and le Roex, 2008; Elliott, 2003; Hole et al., 1984). Negative Ce anomalies are observed in all the Tamba granitoids, with (Ce/Ce*) ratios ranging Fig. 14. FM vs. TiO2, Nb, and Ta in the Tamba granitoids compared with adakites drawn from 0.87 to 0.96. The values of Sr and eNd(105 Ma) in the granitoids from the GEOROCK database (http://georoc.mpch-mainz.gwdg.de). T.Imaoka et al./Lithos 184-187(2014)105-131 125 GEOROCK database (Fig. 14; http://georoc.mpch-mainz.gwdg.de). The Nb and Ta contents of an adakitic melt cannot increase by interac- rich amphiboles (Appendix 1). tion of the melt with mantle peridotite during its ascent, because metasomatized peridotite xenoliths have low Nb and Ta contents of 7.3. Tectonic background of episodic magmatism at ca. 105 Ma in SW Japan 3.9-6.4 ppm, and 0.24-0.43 ppm, respectively. Therefore, the slab melts of the Tamba granitoids are considered to be intrinsically enriched 7.3.1. Subducting oceanic plate in Nb and Ta, despite their negative Nb-Ta anomalies in MORB- Before examination of the tectonic background of episodic normalized pattern. The Tamba granitoids are also rich in TiO2 relative magmatism at ca. 105 Ma, we will look back over the pre-Cretaceous to many reported adakites elsewhere (Fig. 14). As described above subduction history of SW Japan. Early-Middle Jurassic shallow marine (Section 4.2), the Tamba granitoids contain TiO2-rich minerals including deposits are sporadically distributed in the Inner Zone of SW Japan, ilmenite, titanite, and sagenitic biotite with needle-like rutile (Fig. 3D; and most of the Jurassic Inner Zone can be assigned to a forearc region. e.g., Shau et al., 1991), which is consistent with the TiO2-rich nature of The extensive Mino-Tamba accretionary complex indicates that tecton- the granitoids. The rutiles are rich in Nb2O3 (up to 0.27 wt.%), whereas ic accretion beneath the Inner Zone lasted throughout the Jurassic. The Nb has not been detected in the titanite or the ilmenite; therefore, local distribution of the earliest Cretaceous Mino-Tamba accretionary the rutiles are considered to be a major carrier of Nb in the granitoids. complex suggests that the subducting oceanic plate existed beneath The host biotites have higher Mg# and contents of TiO2 relative to Late the Inner Zone, at least until that time (Shuto and Otsuka, 2004; Cretaceous biotites of SW Japan (Fig. 6). The TiO2 content of synthetic Wakita, 1988). biotites increase with increasing temperature, as shown in the hydro- The age of the subducting oceanic plate can be determined approx- thermal experiments of Robert (1976) and Tronnes et al. (1985). The ex- imately from the differences in ages of fossils in the lowermost part of perimental results are consistent with most natural biotites from the pelagic sediments just above the oceanic basement and the terrige- igneous and metamorphic rocks (e.g., Dymek, 1983; Rimsaite, 1964; nous sediments. Oceanic plate subducted during the Jurassic and Early Velde, 1969). The Tamba granitoids containing TiO2-rich biotite are Cretaceous in SW Japan had been in existence for more than 100 my thus considered to have formed at higher temperatures than the Late (Fig. 15). Subduction of the oceanic plate from the latest Early Creta- Cretaceous granitoids of SW Japan. ceous (Albian?） to the Early Miocene under SW Japan produced the Rutile-saturated TiO2 solubility in the slab melt primarily increases Shimanto accretionary complex. An oceanic plate born during or just before the Tithonian had been subducted in Shikoku during the content in the melt, and decreases with pressure (e.g., Hayden and Turonian-Santonian (Ishida, 1998; Nakaseko and Nishimura, 1981), and the age of this plate in the Turonian-Santonian time is estimated ences therein). Xiong et al. (2009) proposed a thermometer taking the to have been ca. 60 Ma (Fig. 15). After that, the age decreases rapidly to- effects of pressure and HzO into account. If we assume the H2O content wards the latest Cretaceous to almost zero (Kiminami et al., 1994; to be 10 wt.%, the Ti02 content of the rutile to be 98 wt.% (Xiong et al., Osozawa, 1992; Taira et al., 1989). Considering the general trend of the 2009), and the pressure 2.0-2.2 GPa (Rapp et al., 1991; Tsuchiya and age of the subducting plate during the Jurassic and Cretaceous in SW Kanisawa, 1994), most of the calculated temperatures for the Tamba Japan, the age of the plate at ca. 105 Ma, when the Kyoto lamprophyres and the adakitic Tamba granitoids formed, is presumed to be more than is assumed to be 15%, the calculated temperatures lie mostly in the 60 Ma, although there are no data on the age of the plate during the mid- o l s m a o Cretaceous. The presumed age of the subducting plate at ca. 105 Ma is fect on temperature. too old to have produced slab melting, because the subduction of very Adakites are frequently associated with various Nb-rich basaltic rocks young (<25 Ma) oceanic plate is required for such melting (Peacock (HNB and NEB; e.g., Castillo, 2012 and references therein), although the et al., 1994). This implies that some factor other than the age of the association is not ubiquitous (Macpherson et al., 2010). During this subducting slab is required to explain the sources of the thermal anom- study, we found lamprophyre dikes with HNB and NEB compositions, aly and the melt production of the adakitic Tamba granitoids. all of which were dated at ca.105 Ma, and are closely associated with the adakites in time and space. Therefore the Tamba granitoids 7.3.2. Slab rollback and upwelling hot asthenosphere Jurassic-Cretaceous igneous rocks occur extensively in eastern one type plays a major role in the petrogenesis of the other. Geochemical China and the Korean Peninsula, although magmatism in the Inner evidence precludes the generation of the Tamba granitoids (adakites) by Zone of SW Japan began at ca. 105 Ma. Recent advances in geochronol- fractionation of the coeval Kyoto lamprophyric magma. The former does ogy with more precise age determinations are revealing temporal and not exhibit fractional crystallization trends consistent with the latter on spatial variations of the Jurassic-Cretaceous magmatism in these re- Harker diagrams (Fig. 7), and the granitoids are distinguished by gions (e.g, Kee et al., 2010; Pei et al., 2011; Sagong et al., 2005; Wang their relatively high values of AlzO3, and approximately constant et al., 1998; Wu et al.,2005a, 2005b, 2007,2011).Some authors have in- contents of Th, Rb and Zr. Furthermore, the initial Sr-Nd isotope voked changes in the dip angle of the subducting slab to explain the data are quite different for the two types. We propose, therefore, spatio-temporal variations of magmatism in eastern China, the Korean that the close spatial and temporal associations of the two types Peninsula, and SW Japan. Sagong et al. (2005) ascribed a Late Jurassic- imply another kind of petrogenetic relationship. Several studies Early Cretaceous lull in magmatism to shallow-angle subduction. have addressed the spatio-temporal associations of Nb-enriched ba- Kiminami et al. (2009), Kiminami and Imaoka (2013) and Zhang et al. salts and adakites in arcs (e.g., Defant et al., 1991; Sajona et al., 1993, (2010) attributed an inland propagation of magmatism during the 1994, 1996; Yogodzinski et al., 1995, 2001). Taking those studies and Early-Middle Jurassic to rapid shallowing of the subduction angle, our data into account, we suggest the following scenario: (1) TiO2-, a hiatus in magmatism during the Late Jurassic to early-Early Creta- Nb-, and Ta-rich adakitic liquids form by slab melting at high tem- ceous to flat-slab subduction, and the trenchward migration of peratures (920-970 °C) during subduction of oceanic lithosphere; magmatism during the Early Cretaceous to slab rollback (Fig. 16A). (2) these liquids may react with sub-arc mantle wedge peridotite; Some authors suggested that slab rollback and the resultant upwell- (3) TiO2-, Nb-, and Ta-rich adakitic liquids could be an effective ing of the asthenosphere might have led to the Early Cretaceous metasomatic agent, forming amphiboles and ilmenites that scavenge magmatism in eastern China (Davis et al., 2001; Li et al., 2003; Yang HFSEs from the liquid; and (4) as suggested by Kepezhinskas et al. et al., 2007; Zhang et al., 2009, 2010). Kiminami and Imaoka (2013) (1996), induced convection in the sub-arc mantle drags the concluded that Cretaceous magmatism in the Inner Zone of SW Japan metasomatized peridotite to depths where it melts, generating Nb- was initiated during the final stage (Albian) of slab rollback. 126 T. Imaoka et al./Lithos 184-187(2014)105-131 Cretaceous Shimanto Belt Ma Mino-Tamba Belt Southern Chichibu Belt [Shikoku] 65.5 ate Early 145.5 Late Middle Early 199.6 Late Triassic Mid. 251.0 Middle Permian Early 299.0 nufer sandstone and mudstone chert andsiliceous shale limestone greenstone 359.2 Fig.15.Oceanic plate stratigraphy for accretionary complexes in the Mino-Tamba Belt (modified from Nakae, 2000),the Southern Chichibu Belt(modified from Matsuoka et al.,1998),and the Cretaceous Shimanto Belt (modified from Taira et al, 1989). As a result of slab rollback, the resultant slab gap was filled with up- Islands (104 Ma; Kanisawa et al., 1983). This event has also been docu- welling hot asthenosphere, producing a transient thermal event and mented from the NE Japan arc, as for example in the Early Cretaceous adakitic and lamprophyric magmatism at ca. 105 Ma. Nakada and high-Sr high-Nb andesites, high-Ti andesites and shoshonites, and Takeda (1995) also proposed a model of hot mantle diapiric upwelling high-Sr andesites similar to bajaites (Tsuchiya et al., 1999). The Early and related convective coupling between the lower crust and upper Cretaceous magmatism began with adakitic and lamprophyric mag- mantle to explain the Early Cretaceous geological events in SW Japan. matic activity. Some of the volcanic rocks in sW Japan, produced con- This anomalous thermal event, the result of unusual tectonic conditions, currently with the lamprophyres, occur within sedimentary basins was not restricted to the Kinki district, but has also been recorded wide- such as those containing the Kanmon and Sasayama Groups. These ly in the SW Japan arc; examples are the Early Cretaceous andesitic basins could have formed in an extensional tectonic regime (Nakada rocks from the Chubu district (107, 106, 102 Ma; Yamada, 2001), and Takeda, 1995). Slab rollback is presumed to lead to an extensional Early Cretaceous HMAs from the Kanmon Group at the westernmost regime in the upper plate (e.g, Biryol et al., 2011; Dilek and end of Honshu (107-103 Ma; Imaoka et al., 1989, 1993; Matsuura, Altunkaynak, 2009; Guillaume et al., 2009). Therefore, it is possible to 1998) and from Kyushu (99 Ma; Kamei et al., 2004), and lamprophyres attribute an extensional regime in the late Early Cretaceous to slab (camptonite and spessartite) from Amami-Oshima in the Ryukyu rollback. T.Imaoka et al./Lithos184-187(2014)105-131 127 Slab rollback during the Early Cretaceous in eastern Asia could if the slab was old (e.g, Calmus et al., 2003; Yogodzinski et al., 2001). In have resulted in elevated geothermal gradients in the mantle wedge, this scenario, the adakitic Tamba granitoids would have resulted from because it produces upwelling of the hot asthenosphere (Kincaid and melting of the subducted slab. The slab-derived adakitic melt, along Griffiths, 2003). Based on analog experiments, Kincaid and Griffiths with fluids liberated from the slab, would then have metasomatized (2003） suggested that upwelling hot asthenosphere resulting from the mantle wedge during ascent, triggered partial melting to generate slab rollback had the potential for slab melting. The geotectonic setting the Kyoto lamprophyres (Fig. 16C). of the Inner Zone of SW Japan is presumed to have been one of subduc- We propose, therefore, that beneath the Kinki district in SW tion-accretion up to the earliest Cretaceous. This suggests that the lith- Japan at around 105 Ma, a transient thermal event was induced by osphere in the Inner Zone around 105 Ma was mostly made up of slab rollback and the consequential upwelling of the hot asthenosphere. tectonically accreted materials, and that the lithospheric mantle was The resultant slab melting produced the adakitic Tamba granitoids, very thin or almost absent (Fig. 16). This may have precluded partial p e pnd m s melting of thickened lower crust origin of adakites (e.g., Atherton and partial melting of the overlying mantle wedge that was metasomatized Petford, 1993), as demonstrated by Kiji et al. (2000). Therefore, the gen- by the slab-derived adakitic Tamba melts. These processes took eration of heat pulses by upwelling hot asthenosphere was able to trig- ger partial melting of the oceanic crust (producing adakitic melts), even 70 km. A) 130 Ma Tan-Lu Fault ↑ extensional tectonics A- & I-type granite ACC & adakitic granite MF migration of MF SSC trench ? +subducting sla.'.- !T++++. +++士 upwelling 十十 →Rollback NW SE B)105 Ma Kyoto alkali basalt ← extensional tectonics trench lithospheric mantle Fig. 16C asthenosphere C) 105 Ma Kyoto ontinentalcrus accretionarycomplex reactionof adakite 口X "'melt with peridotite! solidus... 沙 metasomatized wedge mantle + + H + Tamba granitoids oceanic crust' underplated basaltic rock ++++ +·+ 十 +++· dehydration & sediment melt Fig.16.Schematic cross-sections (A-B) along a SE-NW trend from the ancient trench of SW Japan to the Tan-Lu Fault during the Early Cretaceous showing temporal changes in subducting slab geometry, tectonic regime, and magmatic episodes, and an enlarged cross-section (C) at ca. 105 Ma showing a suggested model for the asthenospheric upwelling and the resulting adakitic and lamprophyric magmatism that formed the Tamba granitoids and the Kyoto lamprophyres. Fig. 16A is modifed from Kiminami and Imaoka (2013). The paleogeographic position of SW Japan is a reconstruction based on Itoh et al. (2006). ACC = accretionary complexes, MF = magmatic front, SSC = Sambosan seamount chain, H = high-(La/Yb)n lamprophyres, L = low-(La/Yb)n lamprophyres, and X = peridotite xenolith. See text for details. 128 T.Imaoka et al./Lithos 184-187(2014)105-131 Acknowledgments Calmus, T, Aguillon-Robles, A, Maury, RC, Bellon, H,Benoit, M., Cotton,J, Bourgois, J, Michaud, F., 2003. Spatial and temporal evolution of basalts and magnesian andesites ("bajaites") from Baja California, Mexico: the role of slab melts. Lithos 66, 77-105. We express our sincere gratitude to Kazuo Kiminami for his kind ad- Castillo, P.R., 2008. Origin of the adakite-high-Nb basalt association and its implications vice and valuable suggestions on tectonics. The authors are also grateful for postsubduction magmatism in Baja California, Mexico. Geological Society of America Bulletin 120, 451-462. to Nobutaka Tsuchiya for many helpful criticisms of the manuscript, and Castillo, P.R, 2012. Adakite petrogenesis. Lithos 134-135, 304-316. to Barry P. Roser for his comments on a draft of our paper. We thank Class, C., le Roex, A.P.,2008. Ce anomalies in Gough island lavas-trace element character- Toshinori Okada for help with K-Ar dating, and Masaaki Owada and istics of a recycled sediment component. Earth and Planetary Science Letters 265, Ami Takashima for XRF analyses. CIPW norm calculations were per- 475-486. Czamanske, G.K,Ishihara, S., Atkin, S.A, 1981. Chemistry of rock-forming minerals of the formed by Surendra P. Verma, using the program SINCLAS. We also Cretaceous-Paleocene batholiths in southwestern Japan and implications for magma thank Nelson Eby for his editorial handling, and Esteban Gazel and an genesis. Journal of Geophysical Research 86, 10431-10469. 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Imaokka 2014 Magmatism 105Ma Kinki SW Japan.txt
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The Island Arc (2005) 14, 582–598Blackwell Science, LtdOxford, UKIAR The Island Arc 1038-48712005 Blackwell Publishing Asia Pty LtdDecember 2005144582598Thematic Article Subducted remnant arc in Cretaceous JapanH. Ueda and S. Miyashita
*Correspondence.
Received 9 May 2005; accepted for publication 5 September 2005.
© 2005 Blackwell Publishing Asia Pty LtdThematic Article
Tectonic accretion of a subducted intraoceanic remnant arc in Cretaceous
Hokkaido, Japan, and implications for evolution of the Pacific northwest
HAYATO UEDA1,* AND SUMIO MIYASHITA2
1Division of Earth and Planetary Sciences, Graduate School of Science, Hokkaido University, North 10 West 8
Kita-ku, Sapporo 060-0810, Japan (email: ueta@mvf.biglobe.ne.jp) and 2Department of Geology, Faculty of
Science, Niigata University, 8050 Ikarashi 2-no-cho, Niigata 950-2181, Japan
Abstract An accretionary complex, which contains fragments of a remnant island arc, was
newly recognized in the Cretaceous accretionary terranes in Hokkaido, Japan. It consistsof volcanics, volcanic conglomerate, intermediate to ultramafic intrusive rocks with island-arc affinity including boninitic rocks, accompanied by chert and deformed terrigenousturbidites. Compared with the results of modern oceanic surveys, the preserved sequencefrom island-arc volcanics to chert, via reworked volcanics, is indicative of intraoceanicremnant arc, because the sequence suggests an inactive arc isolated within a pelagicenvironment before its accretion. The age of a subducting oceanic crust can be discontin-uous before and after a remnant-arc subduction, resulting in abrupt changes in accretionstyle and metamorphism, as seen in Cretaceous Hokkaido. Subduction of such an intra-oceanic remnant arc suggests that the subducted oceanic plate in the Cretaceous was notan extensive oceanic plate like the Izanagi and/or Kula Plates as previously believed bymany authors, but a marginal basin plate having an arc–back-arc system like the present-day Philippine Sea Plate.
Key words: accretionary complex, circum-Pacific ophiolite, Hokkaido, Idonnappu zone,
intraoceanic remnant arc, subduction zone tectonics.
INTRODUCTION
Ophiolitic materials and metavolcanics in ancient
accretionary complexes are the direct records ofalready consumed oceanic plates. Accreted frag-ments of seamounts, oceanic plateaus, abyssalbasements and mid-oceanic ridges have been doc-umented in on-land accretionary complexes (e.g.Miyashita & Katsushima 1986; Petterson et al .
1999; Ueda et al . 2000).
Remnant arcs (Karig 1972) are the other major
oceanic edifices on oceanic plates, especially clus-tered in the western Pacific. Some remnant arcs(Kyushu–Palau Ridge and Amami Plateau;Fig. 1A) are currently subducting along con-
vergent margins, and it is thus expected thatremnant-arc subduction was ubiquitous in ancienttimes. Although marginal basin models have been
widely accepted for supra-subduction zone ophio-lites, with continental remnant arcs in some cases(Harper & Wright 1984), fragments of intraoceanicremnant arcs that accreted from subducting oce-anic plates into continental margins have not beensuccessfully identified in on-land subduction com-plexes. Accreted fragments of remnant arcs havealso not been recognized along the modern inner-trench walls beneath which they are subducting.Intraoceanic remnant arcs are thus commonlythought not to be accreted to continental margins,and their contribution to the upper-plate tectonicsand continental growth has not been wellevaluated.
This paper describes the petrology and strati-
graphy of accreted fragments of a remnant arcnewly found in a Cretaceous accretionary complexin central Hokkaido, Japan, and proposes criteriato identify accreted remnant arcs on land. Thesecriteria may help us to improve our understanding
Subducted remnant arc in Cretaceous Japan 583
of subduction zone tectonics and paleogeographic
reconstruction of oceanic plates.
GEOLOGICAL SETTING
Present-day central Hokkaido is situated at the
collision zone between the northeast Japan andKurile Arcs (Kimura 1986; Arita et al . 1998;
Fig. 1A,B). The northeast Japan Arc terrane com-
prises the Oshima Belt (Jurassic accretionarycomplex and Lower Cretaceous granitoids andvolcanic rocks), the Sorachi–Y ezo Belt (Creta-ceous forearc basin and accretionary complex)and the Hidaka Belt (Cretaceous–Paleogeneaccretionary complex). The Kurile Arc terranecomprises the Nemuro Belt (Cretaceous arc–forearc deposits) and the Tokoro Belt (Creta-ceous–Paleogene accretionary complex). Beforethe marginal basins of the Japan Sea and theKurile Basin opened in the Miocene, the northeast
Japan Arc was part of the Eurasian active conti-nental margin with west-dipping subduction ofthe oceanic plate. Cretaceous accretionary com-plexes, which suffered very low-grade and high-pressure–low-temperature metamorphism, areexposed in the Idonnappu and Kamuikotan zones,respectively , in the Sorachi–Y ezo Belt (Fig. 1C).The accretionary unit containing remnant arcfragments occurs as a serpentinite-bearing nar-row ( <2 km) belt in the Idonnappu zone, here
called the Oku-Niikappu complex (renamed fromthe Oku-Niikappu Dam complex by Ueda &
Miyashita 2003a,b). It lies between the LowerCretaceous and Upper Cretaceous accretionaryunits (Fig. 1C) and is typically exposed aroundthe Niikappu River.
CONSTITUENT LITHOFACIES
The Oku-Niikappu complex is regarded as an
aggregate of fault-bounded slices and blocks con-sisting of four lithofacies associations (here simplytermed ‘facies’): the turbidite, ultramafic, gabbro–diabase and metavolcanic–chert facies (Fig. 2).
The turbidite facies is composed of broken alter-
nation of black mudstone and siliciclastic sand-stone (graywacke), partly containing meter-sizedblocks of chert and metavolcanic rocks. Sandstoneconsists mainly of mineral particles of quartz, pla-gioclase and K-feldspar, and rock fragments of fel-sic volcanic rocks and granitoids. It also containsminor clasts of mudstone, chert, siliceous schist,biotite, zircon, garnet and allanite. This clasticcomposition, suggesting sialic sources, is easilydistinguishable from volcanic sandstone in themetavolcanic–chert facies as described later. Theturbidite facies correlates to eastern parts ofunit O5 of Kiyokawa (1992), who reported mid-Cretaceous (late Albian–Cenomanian) radiolarianfossils from mudstone.
The ultramafic facies consists mainly of massive
and foliated serpentinite. The foliated serpentiniteFig. 1 (A) Present-day tectonic setting around the Hokkaido (Japan) and Philippine Sea regions. AP , Amami Plateau; DR, Daito Ridge; IB M, Izu–Bonin–
Mariana Arc; KPR, Kyushu–Palau Ridge; MT, Mariana Trough; NT, Nankai Trough; PVB, Parece Vela Basin; RT, Ryukyu Trench; SB, Shi koku Basin; WMR,
West Mariana Ridge; WPB, West Philippine Basin. Dotted line shows the Quaternary volcanic front by the Philippine Sea Plate sub duction. Also shown
are the ages of the subducting oceanic crust along the Nankai Trough (Okino et al. 1994) and the Ryukyu Trench (Hilde & Lee 1984), and Deep Sea Drilling
Project sites on and around intraoceanic remnant arcs (open circles with numbers). (B) Tectonic division of Hokkaido. Acc., acc retionary. (C) Geological
outline of south central Hokkaido. Shaded areas indicate the Kamuikotan zone (Km) and the Idonnappu zone (Id). 1, Cretaceous fo rearc basin deposits;
2, Jurassic metabasalt; 3, Lower Cretaceous accretionary complex; 4, Upper Cretaceous accretionary complex; 5, ultramafic rocks; 6, Paleogene
accretionary complex; 7, Cretaceous–Tertiary metaophiolite; 8, Tertiary gneiss and plutonic rocks.
1
2
3
4
56
7
8N
Sorach i-Yezo
BeltHidaka
Belt42°40'N
142°10'E
KmFig. 2
Id
Urakawa
50 kmC B
Oshima Belt
(Jurassic
acc. complex
& Cretaceous arc)Hidaka Belt
(Paleogene acc. complex)
Sorachi-Yezo Belt
(Cretaceous forearc basin & accretionary complex) Tokoro Belt
(Cretaceous
-Paleogene
acc. complex)
Nemuro Belt
(Cretaceous -Paleogene
arc / forearc)
Kurile arc
terrane
NE Japan arc terraneC
Pacific plate
Philippine
Sea plate40°N
30°N
20°N130°E140
°E150
°E
120°E14
0°E
Eurasia (Amur) plateNorth America
(Okhotsk) plate
APNT
RT24-
15 Ma
43->60 MaDR
SB
PVB 448445
451453296
WPB
MTKPR WMRIBMBA
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584 H. Ueda and S. Miyashita
comprises sheared and fully serpentinized
harzburgite, and contains blocks and slices of(olivine–)clinopyroxenite to wehrlite with cumu-late textures, amphibolitized gabbro, metavolca-nics, rodingite and rare olivine–melagabbronoritecumulate. The massive serpentinite comprises acoherent slice originated from clinopyroxene(Cpx)-bearing harzburgite–lherzolite.
The gabbro–diabase facies consists mainly of
metagabbro and mafic to intermediate diabasewith subordinate amounts of diorite, quartz–diorite and tonalite. Most mafic rocks sufferedgreenschist–amphibolite facies metamorphism,overprinted by subgreenschist facies recrystalliza-tion with numerous veins and cataclastic shearbands. These plutonic and hypabyssal rocks com-prise sheeted dyke complex in places (Ueda &Miyashita 2003a).
The metavolcanic–chert facies is the most char-
acteristic lithofacies in the Oku-Niikappu complex.It is composed mainly of metavolcanics, radiolar-ian chert and volcanogenic sedimentary rocks.Beds generally strike northeast–southwest toeast–west, forming a syncline within a slice(Fig. 2).
STRATIGRAPHY IN THE METAVOLCANIC–CHERT
FACIES
Although the metavolcanic–chert facies is inter-
nally cut by minor faults, a near-original strati-graphic succession is observable in several routes(Fig. 3): volcanic rocks in the lower part, alterna-
tion of volcanic conglomerate and chert in the mid-dle part, and chert in the upper part.
The lower part consists mainly of massive lava
and monolithic volcaniclastic rocks (hyaloclastiteand volcanic and pillow breccia), accompanied byrare pillow lava and brown mudstone. They arecomposed mainly of vesicular, aphyric to Cpx-sparsely phyric basalt–andesite. Also seen is greenFig. 2 Geological map and cross-
sections of the Oku-Niikappu complex.Vertical and horizontal scales are in equal
ratio in the cross-sections. Cret.,Cretaceous.
42°40' N90
1 kmXD
Y'Poroshiri LakePelitic &
psammiticsemischistPoroshiriMetaophiolite
Metabasalt &
metagabbro
Y
56Upper
Cret.Lower
Cret.Poroshiri
Meta- ophioliteOku-Niikappu Complex
XX '
Clinopyroxenite–wehrlite
Massive serpentiniteChert
with volcanic sandstone beds
Volcanic conglomerate
with chert beds
Lava and volcaniclastic rocks
Gabbro - diabase faciesTurbidite facies
Metavolcanic-chert facies
Foliated serpentiniteUltramafic faciesY'
Y
1km
Mapped routeOku-Niikappu Complex
Lower Cretaceous
accretionary complex
142 °40' E7040
327580
70
X'Landslide
debris
Landslide
debrisMetabasaltBroken turbidite68
Upper Cretaceous
accretionary complexNiikappu River8070
867070
85AB
C
Fig. 3 Columnar sections of the metavolcanic–chert facies. For local-
ities of sections A–D see Figure 2. F, minor fault.
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Subducted remnant arc in Cretaceous Japan 585
(chromian) Cpx-phyric basaltic andesite, accompa-
nied by pseudomorphs after olivine and picotiteinclusions in cases.
The middle part consists of volcanic conglomer-
ate and sandstone, which intercalate with dark-brown radiolarite and red to white radiolarianchert. The conglomerate beds are poorly sorted,and grade into chert or radiolarite via volcanicsandstone and siltstone (Fig. 4B,C), suggestingtheir sediment gravity flow origin. It comprisesangular–subangular clasts up to boulder size(Fig. 4D), occasionally accompanied by amoe-boidal rip-up clasts of chert. The conglomerateis characteristically polymictic, containing clastsdominantly of volcanic and hypabyssal rocksranging from basaltic to rhyolitic compositions. Italso contains minor clasts of micritic limestone,(olivine–)clinopyroxenite, cumulate and/oramphibolitized gabbro, diorite–tonalite, andscarce serpentinite (Ueda 2003). Volcanic sand-stone contains mineral particles of Cpx, plagio-clase and minor chromian spinel and hornblende,in addition to dominant volcanic fragments. Con-formable succession from monomictic hyaloclas-tite of the lower part to polymictic conglomerateof the middle part is observable in section A ofFigure 3. The frequency and thickness of chert
intercalations increase stratigraphically upwardas the middle part grades into the upper part(Fig. 4A).
The upper part is composed mainly of radiolar-
ian chert. It occasionally intercalates with thin(<1 m thick) graded beds of volcanic sandstone
identical to those in the middle part.
Chert and radiolarite samples from the middle
and upper parts yielded radiolarian fossils(Table 1; Fig. 5). A radiolarite sample R5, whichalternated with volcanic conglomerate in the mid-dle part, contains Pseudodictyomitra carpatica ,
Thanarla pulchra and Mirifusus mediodilatatus .
According to Jud (1994) and Matsuoka (1995),
P. carpatica ranges from latest Tithonian to Bar-
remian, T. pulchra first appears in the late Berri-
asian and M. mediodilatatus disappears in the late
Hauterivian. The age of R5 is thus considered tobe late Berriasian–Hauterivian, and the absence ofCecrops septemporatus implies it is younger than
late Valanginian (KR1 zone of Matsuoka 1995).Sample R4, a rip-up clast of radiolarite in volcanicconglomerate, also contains T. pulchra , indicating
its age is younger than late Berriasian. This sam-ple also yielded Eucyrtidiellum ptyctum , whose
last occurrence is in the early Tithonian (Jud1994). It is interpreted here as a reworked fossil.Other samples contain less diverse assemblageswith poorer preservation, which are not incompat-ible with the age of the former two samples.
PETROLOGY OF IGNEOUS ROCKS
SAMPLES AND ANALYTICAL METHODS
Seventeen selected samples of plutonic and
hypabyssal rocks in the gabbro–diabase facies, andFig. 4 Outcrops of the metavolcanic–
chert facies. (A) Bedded chert (c) of theupper part overlies volcanic conglomer-ate (g) of the middle part. (B) Gradedbeds from volcanic granule conglomer-ate (g) to volcanic sandstone (s) alternatewith a chert bed (c) in the middle part.(C) Alternation of chert beds (c) andgraded beds of volcanic sandstone (s) tobrown mudstone (m) in the middle part.(D) Unsorted and polymictic conglomer-ate in the middle part, consisting of cob-ble- to boulder-sized volcanic clasts.Both angular (a) and subrounded (r)clasts are contained.
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586 H. Ueda and S. Miyashita
lavas and pebbles in the metavolcanic–chert facies
were examined for bulk-rock major and trace ele-ment analysis, performed by Rigaku RIX3000XRF at Niigata University , along with the methodby Takahashi and Shuto (1997). Amygdales, vein-lets and sheared parts were discarded beforepowdering. The results are listed in Table 2.
Other information on magma chemistry was
obtained from the chemical compositions of Cpx,spinel and glass inclusions rarely found in Cpx-phyric hyaloclastite. They were analyzed by JEOLJXA8600SX at Niigata University (accelerationvoltage 15 kV , beam current 13 µA). Beam diame-
ter was broadened to approximately 10 µm for
glass inclusion analyses.BULK-ROCK AND GLASS INCLUSION CHEMISTRY
Although bulk-rock SiO 2 contents scatter, proba-
bly owing to alteration, they show a broad positivecorrelation with immobile and incompatible ele-ments such as Zr (Fig. 6D), and range from basal-tic andesite to andesite in composition. The SiO
2-
saturated nature is confirmed by commonlyquartz-normative glass inclusions (Table 3) ofbasaltic andesite compositions. Both bulk-rock andglass inclusion analyses are divided into at leasttwo groups: the low-Ti (TiO
2 < 1 wt%) and high-Ti
(TiO 2 ≥ 1 wt%) groups (Fig. 6A,E). Glass inclu-
sions of both series belong to the low-K series(Peccerillo & Taylor 1976; Fig. 6C), with K
2O con-Table 1 List of radiolarian fossils from the metavolcanic–chert facies
ID:
Latitude ( °N):
Longitude ( °E):
Horizon:Lithology:Occurrence:R1
42.6715
142.6809
Upper
ch
Ruc in cgR2
42.6714
142.6807
Upper
ch
Alt with ssR3
42.6712
142.6754
Middle
ch
Bedded chR4
42.6709
142.6769
Middle
rad
Ruc in cgR5
42.6706
142.6775
Middle
rad
Alt with cgR6
42.6717
142.6796
Middle
ch
Ruc in ss
Acaeniotyle diaphorogona + cf
Acaeniotyle umbilicata cf +
Acanthocircus trizonalis cf
Alievium helenae cf +
Archaeodictyomitra apiarium ++ +
Archaeodictyomitra excellens cf
Archaeospongoprunum imlayi +
Emiluvia pessagnoi cf
Eucyrtidiellum ptyctum + cf
Gongylothorax favosus cf cf
Hsuum raricostatum aff
Linaresia chrafatensis cf
Mirifusus mediodilatatus +
Obesacapsula cetia cf
Obesacapsula morroensis aff
Pantanellium riedeli + cf
Pantanellium squinaboli cf ++ +
Parvicingula spaerica cf
Parvicingula spinata aff
Pseudodictyomitra carpatica +
Ristola cretacea cf
Sethocapsa kaminogoensis cf aff
Sethocapsa trachyostraca cf
Sethocapsa uterculus ++
Spongocapsula perampla aff
Spongocapsula palmerae ++ cf
Spongocapsula tripes + aff
Syringocapsa agolarium ++ cf
Syringocapsa longitubus aff
Thanarla conica cf cf
Thanarla pulchra ++ cf
Wrangellium puga cf
Xitus spicularius cf
Zhamoidellum ovum +
Zhamoidellum ventricosum +
+, Present; aff, affinis; alt, alternating; cf, confer; cg, volcanic conglomerate; ch, chert; rad, radiolarite; ruc, rip-up clast (amoeboidal
block); ss, volcanic sandstone.
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Subducted remnant arc in Cretaceous Japan 587
tents exclusively higher than those of normal mid-
oceanic ridge basalt (MORB; typically < 0.2 wt%;
Basaltic Volcanism Study Project 1981). Severallow-Ti basaltic andesites (B1–B4) are rich in MgO(8.6–10.9 wt%; Fig. 6B), Cr (Fig. 6H) and Ni con-tents, showing a chemical affinity with boninite,after Le Bas and Streckeisen (1991), although theydo not show any evidence of low-Ca pyroxenes,unlike typical boninites (Crawford et al . 1989).
Relatively undifferentiated (Cr > 200 p.p.m.)
rocks, including boninitic samples, in the low-Tigroup are characterized by very low concentra-tions of high field strength elements (HFSE).Their Y contents (8–17 p.p.m.), significantly lowerthan those of MORB (20–70 p.p.m.; Pearce 1982)
with equivalent Cr contents (200–400 p.p.m.), canbe attributed to extensive partial melting and/orrefractory source peridotite. Whereas the mildlyK
2O-rich character of the low-Ti group is inconsis-
tent with such highly depleted nature, and thussuggests selective enrichment probably by H
2O
fluid. The SiO 2-oversaturated compositions also
prefer the presence of an H 2O component in the
upper mantle (Kushiro 1972). These chemical char-acteristics and the conditions deduced thus sug-gest low-Ti series rocks of island arc origin. HFSEratios such as Zr/Nb (15–36; Fig. 6F) are alsowithin the range of island arc basalts (BasalticFig. 6 Selected major and trace element
plots for bulk-rock and glass inclusioncomposition of igneous rocks. (A–C) TiO
2,
MgO and K 2O variations versus SiO 2,
respectively; (D–H) SiO 2, Ti, Nb, Y and Cr
variations versus Zr, respectively, inascending order. Low-, medium- andhigh-K fields after Peccerillo and Taylor(1976).Zr/Nb = 10 20
30012
50 60 70
SiO (wt%)2TiO (wt%)2
Zr(p.p.m.)0100200300400500
05 0 1 0 0 150Cr (p .p.m.)SiO (wt%)201234
50 52 54 56 58 60K2O (wt%)
low-Kmedium-Khigh-K
02468
05 0 1 0 0 150
Zr(p.p.m.)Nb (p .p.m.)SiO (wt%)2SiO 2(wt%)
506070
05 0 1 0 0 150
Zr(p.p.m.)
01020304050
05 0 1 0 0 1 5 0
Zr (p.p .m.)Y (p.p.m.)051015
50 60 70MgO (wt%)
Glass
inclusion
G1
G2
G3Gabbro
-diabase
facies
High-Ti
Low-Ti
Low-Ti
(Cr-rich)LavaMetavolcanic
-chert facies
Pebble0200040006000800010000
05 0 1 0 0 150
Zr(p.p.m.)Ti (p.p.m.)A B C
D E F
G HFig. 5 Radiolarian fossils from chert and radiolarite of the metavolcanic–chert facies. (1–10) Sample R5 from a radiolarite bed altern ating with volcanic
conglomerate in the middle part (1, Mirifusus mediodilatatus (Rüst); 2, Pseudodictyomitra carpatica (Lozyniak); 3, Archaeodictyomitra apiarium (Rüst);
4, Sethocapsa uterculus (Parona); 5, Pantanellium squinaboli (Tan); 6, Parvicingula cf. sphaerica Steiger; 7, Syringocapsa aff. longitubus Jud; 8, Thanarla
pulchra (Squinabol); 9, Archaeospongoprunum imlayi (Pessagno); 10, Syringocapsa agolarium Foreman). (11, 12) Sample R4 from a radiolarite block
(rip-up clast) in volcanic conglomerate in the middle part (11, Thanarla pulchra (Squinabol); 12, Eucyrtidiellum ptyctum (Riedel & Sanfilippo)).
(13) Spongocapsula palmerae Pessagno from a chert sample R2 in the upper part.
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588 H. Ueda and S. Miyashita
Table 2 Major and trace element composition of igneous rocks
ID: B1 B2 B3 B4 B5 B6 B7†B8 B9 B10 B11 B12† B13 B14 B15 B16 B17
Sample no.: 816-7 816-7B 816-9A 817-18 807-14cgB7 Ni-10 909-7 726-19 816-10 817-19 807-7 909-9 720-6 720-9 720-11 725-20 817-1
Unit: MC(l) MC(l) MC(l) MC(l) MC(m) MC(m) GD MC(l) MC(l) MC(l) MC(m) MC(m) GD GD GD MC(l) MC(l)Occurrence: PL PL VB ML Cg Cg MP ML ML VB Bl Cg MP MP MP ML MLRock type: Cpx-p ba Cpx-p ba Cpx-p ba Ol-Cpx-p ba Ol-Cpx ba Ol-Cpx ba meta-g Pl-s db Cpx-Pl a Opx?-Pl-Cpx d Cpx-s a Cpx-Pl-s ba aph db Hbl? t Qtz-tp Cpx-Pl a Cpx-s ba
Chem. type: LT(HCr) LT(HCr) LT(HCr) LT(HCr) LT(HCr) LT(HCr) LT(HCr) LT LT LT LT LT LT LT LT HT HT
SiO
2 53.82 55.15 56.26 52.18 53.66 58.00 52.92 72.20 62.07 64.03 59.93 55.30 57.83 61.73 67.89 56.69 54.07
TiO 2 0.35 0.38 0.35 0.30 0.71 0.47 0.53 0.53 0.40 0.37 0.71 0.74 0.69 0.79 0.62 1.48 1.45
Al2O3 11.50 13.14 13.14 13.15 14.97 14.47 14.92 12.59 15.44 13.81 17.17 16.73 15.37 14.90 14.90 15.83 14.55
Fe2O3 7.35 7.81 6.17 8.04 9.92 7.77 9.41 4.72 6.42 6.59 6.75 9.74 11.97 7.57 6.39 8.79 11.50
MnO 0.18 0.14 0.13 0.15 0.11 0.13 0.14 0.17 0.27 0.11 0.06 0.11 0.09 0.09 0.08 0.17 0.38MgO 10.72 10.83 8.64 10.91 8.26 5.11 9.69 3.45 5.25 5.23 1.86 5.66 5.21 2.56 1.97 4.79 5.38CaO 11.08 7.81 8.16 8.80 5.47 8.11 7.74 1.58 3.58 4.63 4.38 1.87 2.50 9.88 2.64 4.67 5.03Na
2O 3.52 2.56 1.31 2.29 3.95 3.80 3.27 4.27 6.22 2.49 6.12 2.71 4.17 1.27 5.12 6.55 2.99
K2O 1.12 2.20 5.48 3.25 2.45 1.84 1.22 0.27 0.22 3.38 1.92 6.82 1.22 0.06 0.29 0.42 2.84
P2O5 0.07 0.07 0.05 0.06 0.14 0.08 0.04 0.10 0.06 0.07 0.14 0.16 0.09 0.37 0.12 0.22 0.16
Total 99.71 100.09 99.69 99.13 99.64 99.78 99.88 99.88 99.93 100.71 99.04 99.84 99.14 99.22 100.02 99.61 98.35
LOI 4.55 3.51 3.00 4.18 3.96 4.05 2.75 2.45 4.23 3.23 3.40 3.95 3.04 2.32 2.03 1.30 5.23FeO*/MgO 0.62 0.65 0.64 0.66 1.08 1.37 0.87 1.23 1.10 1.13 3.26 1.55 2.07 2.66 2.92 1.65 1.92
Ba 66 269 475 1232 130 148 164 18 18 364 177 250 96 7.5 23 46 151
Cr 383 367 351 385 402 259 218 18 91 78 63 27 25 17 16 49 23Nb 0.2 1.4 1.0 0.4 0.6 0.6 1.1 2.4 1.1 1.4 1.1 1.3 1.1 1.3 1.7 4.0 2.2Ni 65 73 73 104 59 131 70 2.5 41 21 28 37 13 n.d. n.d. 25.7 5.7Rb 6.5 16 33 51 39 33 14 4.2 1.1 24 31 59 1.5 0.4 1.8 7.2 12Sr 155 159 169 442 207 254 294 153 149 168 363 143 241 416 208 274 122V 190 204 238 180 267 264 332 85 122 189 206 293 257 108 67 284 399
Y1 1 1 0 8.2 9.5 17 15 17 23 10 10 33 37 22 24 33 37 25
Zr 23 24 24 23 52 43 18 114 59 31 72 62 55 58 134 101 71
*Total iron as FeO;
†analyzed at Hokkaido University .
a, andesite; ba, basaltic andesite; Bl, block in conglomerate; Cg, clasts in conglomerate; Cpx, clinopyroxene; d, dacite; db, d iabase; g, gabbro; GD, gabbro–diabase facies; Hbl, hornblende; HCr, high-Cr subtype; HT , high-
Ti series; l, lower part; LOI, loss on ignition; LT , low-Ti series; m, middle part; MC, metavolcanic–chert facies; ML, massive lava; MP , massive pluton; n.d., below detection limit; Ol, olivine; Opx, orthopyroxene; p, porphyritic;
Pl, plagioclase; PL, pillow lava; Qtz, quartz; s, sparsely phyric; t, tonalite; tp, tonalitic porphyry; VB, volcanic breccia.
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Subducted remnant arc in Cretaceous Japan 589
Volcanism Study Project 1981). Cr-poorer, more
differentiated rocks in the low-Ti group showgreater variation and generally higher concentra-tions in HFSE (Fig. 7B). They tend to show nega-tive spikes at Ti in trace-element patterns as Crdecreases, presumably resulting from fraction-ation of Fe–Ti oxides.
The high-Ti group samples show a differentiated
nature poor in Cr and Ni, and rich in HFSE. Theyare also regarded as a variety of island arc rocksin terms of their SiO
2- and K 2O-rich composition.
Although their high HFSE concentrations with flatpatterns (Fig. 7C) are similar to those of MORB,they may have evolved from HFSE-poorer magmathan MORB, taking their Cr-poor nature intoaccount. HFSE ratios such as Zr/Nb and Zr/Y(Fig. 6F ,G), except those including Ti (Fig. 6E), arealmost identical to those of the low-Ti group. How-ever, Ti does not show negative spikes (Fig. 7C),unlike the differentiated low-Ti rocks, suggestingthat the high-Ti rocks were not significantlyaffected by fractionation of Fe–Ti oxides. Thesecharacteristics imply that both the high- and low-Ti series originated from source mantle regionswith similar HFSE ratios. The distinct differencein TiO
2 trends can be attributed to differing timingof Fe–Ti oxide crystallization, presumably con-
trolled by oxidation state during differentiation.
CLINOPYROXENE
Relic Cpx were analyzed for clinopyroxenite–wehrlite, gabbro–diabase and volcanic rocks ofboth igneous bodies and pebbles. Several to sev-eral tens of analyses, whose mean values (repre-sentatives) are given in Table 4, were carried outon three or more phenocrysts for each sample.
Two compositional trends, both low-Ti ( <0.015
per formula unit, pfu) and high-Ti series(≥0.015 pfu) are observed in the Ti vs Mg/
(Mg + Fe) value (Mg#) plot (Fig. 8A), correspond-
ing to the bulk-rock low- and high-Ti series,respectively . In the low-Ti series, Ti contents gen-tly increase with decreasing Mg#. The high-Tiseries shows a steeply increasing Ti trend as Mg#decreases to 0.75–0.70, followed by a broad Ti-decreasing trend probably as a result of precipita-tion of Fe–Ti oxides at varying timing.
Clinopyroxenes of both the low- and high-Ti
series plot in the nonalkalic (not shown) and oro-genic fields (i.e. island arc field; Fig. 8B) of Leter-rier et al . (1982). The high-Ti series samples tendTable 3 Representative electron microprobe analyses of glass inclusions in clinopyroxene phenocrysts of hyaloclastite
Sample no.: G1: 725-16 G2: 725-15 G3: 725-19
Point no.: 22 92 142 119 257 210 210 σ
SiO 2 53.99 55.76 54.10 56.44 51.95 55.99 0.20
TiO 2 0.74 0.79 0.79 0.66 1.08 1.01 0.05
Al2O3 15.48 15.88 15.41 15.58 14.65 14.80 0.11
FeO 8.75 8.80 8.66 7.81 10.01 9.48 0.27
MnO 0.24 0.15 0.24 0.10 0.33 0.08 0.04MgO 5.49 4.78 5.23 4.92 6.28 4.34 0.07CaO 8.98 8.29 9.21 8.59 9.78 7.69 0.11Na
2O 2.85 3.24 3.00 3.01 2.77 3.44 0.08
K2O 0.51 0.43 0.54 0.50 0.49 0.87 0.04
P2O5 n.d. n.d. 0.03 0.04 0.04 0.11 0.03
Total 97.03 98.12 97.21 97.65 97.38 97.81
FeO/MgO 1.59 1.84 1.66 1.59 1.59 2.18
CIPW norm†
qz 5.8 7.6 5.4 9.5 2.0 7.2
or 3.0 2.6 3.2 3.0 2.9 5.2ab 24.1 27.4 25.4 25.4 23.4 29.1an 27.9 27.5 27.0 27.5 26.1 22.3ne ––––––
ol ––––––
hy 19.8 19.0 18.1 17.5 21.1 17.8di 13.7 11.3 15.3 12.2 18.3 12.6il 1.4 1.5 1.5 1.2 2.0 1.9mt 1.4 1.4 1.4 1.3 1.6 1.5ap – – 0.1 0.1 0.1 0.3
†Fe3+/Fetotal is assumed to be 0.1.
–, Absent; n.d., below detection limit.
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590 H. Ueda and S. Miyashita
to plot in the overlapping field of orogenic and non-
orogenic Cpx reflecting their Ti-rich nature. Manyof the low-Ti samples are exclusively in the oro-genic field, whereas some Cr-rich ones (in Cpx-phyric lava and diabase) plot in the overlappingfield.
SPINEL
Compositions of spinels (picotites) in Cpx-phyriclava and pebbles and clinopyroxenite–wehrlite (allbelonging to low-Ti series), and of coarse-graineddetrital spinels (presumably derived from peridot-
ite) in volcanic sandstone were analyzed (Table 5).
Spinels in Cpx-phyric lava and pebble samples
(B4–6) show Cr/(Cr + Al) values (Cr#) ranging
from 0.79 to 0.86, higher than the compositionalrange of ocean floor basalts (Barnes & Roeder 2001;Fig. 9A). Spinels in olivine–clinopyroxenite and
wehrlite show less Cr# (approximately 0.6 and 0.7)than volcanic rocks. Spinels of both volcanic rocksand clinopyroxenite–wehrlite are also character-ized by low concentrations of Ti (Fig. 9B), plottingwithin the island arc to boninite field by Arai(1992). The Cr# of detrital spinels ranges from0.72 to 0.93, and is significantly higher than thecompositional range of abyssal peridotite, suggest-ing a more refractory island arc mantle (Fig. 9A).
The very high Cr# values, exceeding 0.8 for
Cpx-phyric lavas, are consistent with their boni-nitic bulk-rock characteristics. These magmas,which probably resulted from low-pressure,hydrous melting of depleted peridotite (Crawfordet al . 1989), are expected to have left highly refrac-
tory residual peridotite, which inferably derivedthe detrital spinels also having very high Cr#.
DISCUSSION
REMNANT ARC OPHIOLITE
One of the unique characteristics in the Oku-
Niikappu complex is a conformable successionfrom island-arc type volcanics via volcanic con-glomerate to chert within the metavolcanic–chertfacies (Fig. 3). Because petrologic characteristicsare common to lava/hyaloclastite (lower part) andvolcanic clasts in conglomerate (middle part), thelower part is regarded as part of an island arcvolcanic edifice, and the middle part conglomerateas their reworked debrites. The gabbro-diabaseand ultramafic facies must also have comprised thesame intraoceanic arc, because clasts of plutonicand ultramafic rocks having common petrologic
Fig. 7 Bulk-rock trace-element patterns for igneous rocks. Mid-oce-
anic ridge basalt-normalizer after Sun and McDonough (1989) with Crafter Pearce (1982). (A) Low-Ti group (Cr > 200 p.p.m.); (B) low-Ti group
(Cr < 200 p.p.m.); (C) high-Ti group. Also plotted in (C) are reference
patterns of high-Ca boninite (Pearce et al. 1992), island arc tholeiite and
calc-alkali basalt (Pearce 1982).0.1110100
Sr K Rb B aN b P Z r Ti Y C rROCK / MORB
0.1110100
Sr K Rb Ba Nb P Zr Ti Y C rROCK / MORB
0.1110100
Sr K Rb Ba Nb P Zr Ti Y C rROCK / MORBIsland arc tholeiite
High-Ca boniniteCalc-alkali basalt(A)
(B)
(C)
Fig. 8 (A) Ti vs Mg# and (B) Ti + Cr vs
Ca plots for clinopyroxenes. Mean valuesof individual samples are shown as sym-bols, and all analyses as gray dots. Dis-criminant fields in (B) are after Letterieret al. (1982).
00.010.020.03
0.6 0.7 0.8 0.900.020.040.06
0.6 0.7 0. 80 . 9 1Gabbro-diabase faices
Lower part (lava & hyaloclastite)
Middle part (conglomerate)Metavolcanic-chert faciesUltramafic facies
(clinopyroxenite-wehrlite)Mean valuesAll analyses
Clinopyroxenite clastsGabbro clastsLava clast
Low-Ti trendHigh-Titrend
Mg# (Mg/Mg+Fe)TiTi
CaOrogenicTi + CrNon-orogenic(A) (B)
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Subducted remnant arc in Cretaceous Japan 591Table 4 Compositions (mean values) of clinopyroxene in selected samples
ID: P1 (B5) P2 (B6) P3 P4 P5 P6 P7 (B4) P8 (G1) P9 (G2) P10 (G3) P11 (B16) P12 P13 P14 P15 P16 P17
Sample: 80714cgB7 Ni-10 80714pxB 80714cgA1 808–9B 80714spB3 817–18 725–15 725–16 725–19 725–20 1009–5 1009–18 721–9 90906 1009–19B 817–6BUnit: MC(m) MC(m) MC(m) MC(m) MC(m) MC(m) MC(l) MC(l) MC(l) MC(l) MC(l) GD GD GD GD GD UMLithology: Ol-Cpx ba Ol-Cpx ba cpxnite Cpx-cum Pl-Cpx ba Cpx ba Ol-Cpx-p ba Cpx HC Cpx HC Cpx HC Cpx-Pl a Cpx-p db db Ol-gn db db WehrliteType: LT(HCr) LT(HCr) LT LT LT HT LT(HCr) LT LT LT & HT? HT LT(Cpx-p) LT LT F HT LT
Occurrence: Clast Clast Clast Clast Clast Clast ML HC HC HC ML SD SD Block MP SD UMn:2 9 9 1 3 4 2 58 53 6 4 91 2 1 2 0 4 8 42 72 20 54 41 58 47
SiO
2 52.73 52.98 52.64 52.26 51.32 52.02 53.32 52.71 52.83 52.05 50.63 53.05 52.62 50.92 51.00 50.98 52.47
TiO 2 0.25 0.09 0.15 0.16 0.42 0.63 0.12 0.26 0.23 0.31 0.77 0.11 0.09 0.33 0.46 0.66 0.18
Al2O3 2.20 2.23 2.03 1.92 2.08 3.06 2.18 2.30 2.32 2.32 3.36 1.65 0.85 2.04 3.07 2.46 2.26
FeO 5.45 5.47 4.28 5.33 10.70 7.22 4.45 5.92 5.54 5.97 9.43 4.59 7.14 8.49 7.12 10.43 3.78
MnO 0.15 0.03 0.15 0.04 0.33 0.19 0.13 0.13 0.17 0.19 0.25 0.13 0.04 0.30 0.10 0.30 0.10MgO 17.23 18.23 16.73 16.25 16.13 16.30 17.71 17.70 17.81 17.71 15.87 18.04 14.66 14.97 16.11 15.67 17.28CaO 21.62 20.32 23.03 22.87 17.82 21.21 22.64 21.07 19.94 21.06 19.49 22.22 23.82 22.11 21.32 19.50 22.92Na
2O 0.23 0.20 0.31 0.27 0.30 0.23 0.12 0.21 0.23 0.20 0.29 0.16 0.47 0.30 0.26 0.27 0.21
K2O 0.00 0.00 0.00 0.00 0.03 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0.00
Cr 2O3 0.30 0.54 0.47 0.36 0.01 0.06 0.24 0.42 0.39 0.37 0.15 0.37 0.06 0.01 0.20 0.08 0.54
Total 100.16 100.09 99.79 99.46 99.14 100.92 100.91 100.72 99.46 100.20 100.24 100.32 99.75 99.47 99.64 100.35 99.74
Si 1.929 1.930 1.931 1.932 1.929 1.903 1.931 1.919 1.936 1.908 1.880 1.933 1.963 1.914 1.893 1.900 1.921
Ti 0.007 0.002 0.004 0.004 0.012 0.017 0.003 0.007 0.006 0.009 0.022 0.003 0.003 0.009 0.013 0.019 0.005Al 0.095 0.096 0.088 0.084 0.092 0.132 0.093 0.099 0.100 0.100 0.147 0.071 0.037 0.090 0.134 0.108 0.098Fe 0.167 0.167 0.131 0.165 0.336 0.221 0.135 0.180 0.170 0.183 0.293 0.140 0.223 0.267 0.221 0.325 0.116
Mn 0.005 0.001 0.005 0.001 0.011 0.006 0.004 0.004 0.005 0.006 0.008 0.004 0.001 0.010 0.003 0.009 0.003Mg 0.940 0.990 0.915 0.896 0.904 0.889 0.956 0.960 0.973 0.968 0.879 0.980 0.815 0.839 0.891 0.871 0.943Ca 0.849 0.794 0.907 0.907 0.718 0.833 0.879 0.823 0.784 0.828 0.777 0.869 0.953 0.892 0.849 0.780 0.900Na 0.016 0.014 0.022 0.019 0.022 0.016 0.008 0.015 0.016 0.014 0.021 0.011 0.034 0.022 0.019 0.020 0.015K 0.000 0.000 0.000 0.000 0.001 0.000 0.000 0.000 0.000 0.001 0.000 0.000 0.000 0.000 0.000 0.000 0.000
Cr 0.012 0.022 0.019 0.015 0.000 0.002 0.010 0.017 0.016 0.015 0.006 0.015 0.002 0.000 0.008 0.003 0.022Total 4.020 4.016 4.022 4.023 4.025 4.019 4.019 4.024 4.006 4.032 4.033 4.026 4.031 4.043 4.031 4.035 4.023
Mg# 0.849 0.856 0.875 0.844 0.729 0.801 0.876 0.842 0.851 0.841 0.750 0.875 0.785 0.759 0.801 0.728 0.890
Numbers of cations are on a six-oxygen basis. a, andesite; ba, basaltic andesite; Cpx, clinopyroxene; cpxnite, clinopyroxenite; cum, cumulate gabbro; db, diabase; F , felsic; GD, gabbro–diabase facies; gn, gabbronorite;
HCr, high-Cr subtype; HC, hyaloclastite; HT , high-Ti series; l, lower part; LT , low-Ti series; m, middle part; MC, metavolcanic –chert facies; ML, massive lava; MP , massive pluton; n, number of analyses; Ol, olivine; p,
porphyritic; Pl, plagioclase; SD, sheeted dyke; UM, ultramafic facies.
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592 H. Ueda and S. Miyashita
characters to these facies are contained in the con-
glomerate. The interbedded and overlying radi-olarian chert, and the absence of siliciclasticintercalation of matured continental crust originsuggest a pelagic, intraoceanic environment. Thesuccession thus suggests that an island arc endedits activity , was then partly eroded and reworkedas debrites, and was finally isolated in a pelagicenvironment. This history correlates to those ofintraoceanic remnant arcs left in pelagic environ-ments after back-arc spreading.According to oceanic surveys on modern rem-
nant arcs, arc-type volcanic and plutonic rocks,rarely accompanied by serpentinite and amphiboleschist, were dredged and cored from remnant arcsand surrounding areas in the Philippine Sea Platesuch as the Amami Plateau, Daito Ridge, Kyushu–Palau Ridge, and eastern foot of the West Mariana
Ridge (Fig. 1A; Yuasa & Watanabe 1977;Tokuyama et al . 1980; Kroenke et al . 1981b; Nat-
land 1982; Shiki et al . 1985). The arc basements are
overlain by debris aprons of reworked coarseTable 5 Representative spinel analyses
Sample/ID:
Occurrence:Unit:90910
Detrital
MC(m)90911
Detrital
MC(m)P1
Ol-Cpx ba
MC(m)P2
Ol-Cpx ba
MC(m)P3
cpxnite
MC(m)P7
Ol-Cpx-p ba
MC(l)P17
Wehrlite
UM
SiO
2 0.06 0.03 0.00 0.10 0.01 0.03 0.02
TiO 2 0.17 0.21 0.61 0.21 0.63 0.34 0.51
Al2O3 14.50 14.75 9.07 11.33 19.56 8.93 16.09
FeO 19.27 22.04 26.49 23.65 37.22 24.03 29.93
MnO 0.49 0.23 0.39 0.29 0.32 0.32 0.27MgO 9.96 12.27 10.64 11.23 7.85 10.96 9.76CaO 0.02 0.04 0.06 0.20 0.29 0.10 0.02Na
2O 0.03 0.00 0.00 0.00 0.00 0.00 0.00
K2O 0.00 0.00 0.02 0.00 0.00 0.00 0.00
Cr 2O3 55.37 47.86 52.90 50.26 30.92 55.28 43.10
Total 99.87 97.43 100.18 97.27 96.80 99.99 99.70
Si 0.002 0.001 0.000 0.003 0.000 0.001 0.001
Ti 0.003 0.004 0.013 0.004 0.014 0.007 0.011Al 0.456 0.484 0.296 0.376 0.682 0.289 0.530Fe
3+0.000 0.022 0.050 0.033 0.270 0.003 0.103
Fe2+0.430 0.490 0.562 0.523 0.651 0.548 0.597
Mn 0.011 0.005 0.009 0.007 0.008 0.008 0.006Mg 0.396 0.509 0.438 0.471 0.346 0.449 0.407Ca 0.000 0.001 0.002 0.006 0.009 0.003 0.001Na 0.001 0.000 0.000 0.000 0.000 0.000 0.000K 0.000 0.000 0.001 0.000 0.000 0.000 0.000
Cr 1.646 1.484 1.629 1.577 1.020 1.692 1.344Total 2.945 3.000 3.000 3.000 3.000 3.000 3.000
Mg# 0.479 0.510 0.438 0.474 0.347 0.450 0.405Cr# 0.783 0.754 0.846 0.807 0.599 0.854 0.717
Ferrous and ferric iron is distributed assuming three total cations on a four-oxygen basis. ba, basaltic andesite; Cpx, clinopy roxene;
cpxnite, clinopyroxenite; l, lower part; m, middle part; MC, metavolcanic-chert facies; Ol, olivine; p, porphyritic; UM, ultram afic facies.Fig. 9 Spinel composition for volcanic
and plutonic rocks. Compositional fieldsare after Barnes and Roeder (2001) in (A),and after Arai (1992) in (B). 0123
00 . 2 0 . 4 0 . 6 0 . 8 1
Cr/(Cr+Al)TiO 2(wt%)
00.20.40.60.81
0 0.2 0.4 0.6 0.81
Mg/(Mg+Fe2+)Cr/(Cr+Al)Intra-plate
MORB
ArcBoniniteCpx-phyric lava
Cpx-phyric volcanic clastOl-clinopyroxeniteOl-clinopyroxenite clast Detrital mineral particleAB
Abyssal peridotiteOcean floor basaltIsland arc
tholeiiteBoninite
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Subducted remnant arc in Cretaceous Japan 593
deposits, with frequent intercalations of pelagic
biogenic sediments in places, and are finally cov-ered by pelagic sediments (Fig. 10; Ingle et al .
1975; deVries Klein et al . 1980; Kroenke et al .
1981a,b). These sequences may represent initialarc magmatism, syn- to postrifting proximal dep-osition, and subsequent isolation with pelagic dis-tal sedimentation, in ascending order. Therefore,the ‘arc–debrite–pelagic succession’ (Fig. 10), canbe diagnostic of ancient intraoceanic remnant arcincorporated in orogenic belts. The arc basementmay also be metamorphic or intrusive rocks in casedeeper rocks had been exhumed. Gabbroic base-ment overlain by a fining-upward debrite sequenceon the eastern foot of the West Mariana Ridge isone of the modern examples, and crustal attenua-tion during the initial back-arc rifting has beenproposed for the exhumation of middle to lowercrustal rocks (Natland 1982). Metagabbro andultramafic clasts and detrital spinel grains in deb-rites of the Oku-Niikappu complex may have alsoderived from middle to lower crust and mantlerocks (i.e. gabbro–diabase and ultramafic facies) ofthe arc extensionally unroofed during the back-arcrifting (Fig. 10).
The interbedded occurrence of arc-derived
debrites and pelagic sediments also suggests aremnant-arc origin, even if the original succes-sion was broken. They differ from trench-filldeposits, usually overlying (i.e. younger than)pelagic sediments (Matsuda & Isozaki 1991). Theremnant-arc debrites, consisting dominantly ofbasaltic to andesitic clasts from immature intra-oceanic arcs, are also petrographically distin-guishable from usually siliciclastic trench-fillturbidites (graywacke) derived from sialic conti-nental margins. The turbidite facies of theOku-Niikappu complex is thus interpreted astrench-fill turbidites in terms of its siliciclasticcomposition and its younger age than the
metavolcanic–chert facies.
IMPLICATION FOR PALEOGEOGRAPHY OF OCEANIC PLATE
Identification of remnant arcs gives constraints onpaleogeographic reconstruction of already con-sumed oceanic plates. The Oku-Niikappu Rem-nant Arc in the Idonnappu zone accretionarycomplex indicates that the subducted oceanic platein the Cretaceous had back-arc basin(s) inside,and had an intraoceanic arc somewhere on itsmargin. This deduced marginal basin plate, simi-lar to the present-day Philippine Sea Plate, isinconsistent with the previous models, which haveassumed subduction of huge oceanic plates(Izanagi, Kula and Pacific Plates) along the Eur-asian margin (Engebretson et al . 1985; Maruyama
& Seno 1986; Kimura et al . 1994), but similar to
those assuming another oceanic plate with anintraoceanic arc between Eurasia and the hugeoceanic plates (Stavsky et al . 1990; Zonenshain
et al . 1990).
Figure 11 shows the reconstructed temporal
change of age of the subducted oceanic plate(s) incentral to southwest Hokkaido. According to theconcept of ocean plate stratigraphy (Matsuda &Isozaki 1991), the oldest age of pelagic sedimen-tary rocks or volcanic basement rocks gives theyoungest age limit of the subducted oceanic crust,and the age of continent-derived clastic rocks sug-gests the timing of initial subduction at the conti-nental margin (trench) in a single accretionaryunit. The age difference between the oldest oce-anic rocks and trench-fill sediments can thus beapproximated to be the age of the oceanic crust atthe timing of its subduction. The Oku-NiikappuRemnant Arc crust, which formed until the earli-est Cretaceous (the age of the chert and radiolariteFig. 10 Stratigraphic comparison of
the metavolcanic–chert facies of theOku-Niikappu complex with intraoceanicremnant arcs in the present PhilippineSea Plate (Deep Sea Drilling ProjectSite 445, deVries Klein et al. 1980;
Site 448, Kroenke et al. 1981b; Site 296,
Ingle et al. 1975; Site 451, Kroenke et al.
1981a).
Middle Eocene
Middle
Oligoc.Oligocene Miocene Pliocene
Late Miocene P.
Mid-Cretaceous
(Albian - Cenomanian)
100mOku-Niikappu
Complexcalcareous ooze,chalk & limestoneSiliciclastic turbidite
chert & radiolariteTuff, lapillituff,
& volcanic brecciaVolcanic sandstone
Volc. conglomerate
Arc lava clay & silt
1. Volcanic basement
(parent arc activity)2. Syn- to post- volcanic
coarse clastics
(intra-arc to back-arc rifting) 3. Biogenic deposits
(pelagic isolation after
back-arc spreading) 4. Continent-derived clastics
(subduction along the
continental margin)
ridge crestMioc.DSDP 448Kyushu-Palau
Ridge
west slopeMioc.DSDP 296West Mariana
Ridge
edge of
ridge crest
DSDP 451Daito Ridge
graben
DSDP 445
Earliest Cretaceous
(Berriasian - V alanginian)Late Oligocene
Parent intraoceanic arcDaughter arcRemnant arcremnant arc ophiolite
(accretionary complex)
Back-arc basin crust
PeridotiteSerpentinite
New back-arc basin
Exhumation of serpentinite
& plutonic rocksPelagic coverTrench-fill turbidite
Continental
Syn-rift debritemarginLate E.
Late Oligocene
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594 H. Ueda and S. Miyashita
covering volcanics), accreted to the Eurasian con-
tinental margin in the mid-Cretaceous (the age ofthe turbidite facies), and the age difference is thusestimated as 30–40 my (Fig. 11).
During Jurassic–Early Cretaceous time, the
subducted oceanic plate was as old (i.e. had agedifferences as large) as 150–200 my , whereas it wasless than 40 my old in mid-Cretaceous–Paleogenetime in central Hokkaido. There is an age gap ofthe subducted oceanic crust in mid-Cretaceoustime in accordance with the accretion timing ofthe Oku-Niikappu Remnant Arc. This temporalchange records the geographic age distribution ofthe subducted oceanic plate (Fig. 11). The deducedage gap across the remnant arc leads to the follow-ing scenario: a parent intraoceanic arc, which wasbuilt on the edge of an old (Carboniferous–Permian) oceanic plate probably in Jurassic time,
suffered back-arc spreading in the earliest Creta-ceous. The old part of the plate was subductedbeneath the Eurasia plate in the Jurassic–EarlyCretaceous, and the newly formed back-arc basinparts subsequently began to subduct since themid-Cretaceous (Fig. 11). Subducted oceanicspreading centers deduced from normal MORBbasalts and dolerite, which erupted in andintruded into trench-fill turbidites in the HidakaBelt (Miyashita & Katsushima 1986; Kiminamiet al . 1999), are interpreted here as those of a well-
developed back-arc basin. Initiation of the parentintraoceanic arc might have resulted from a breakwithin the old oceanic plate probably correlatingto the Izanagi Plate (Engebretson et al . 1985). A
part of presumably the Izanagi Plate was thusseparated to become a new marginal basin plate,beneath which the rest of the Izanagi Plate mayhave subducted. This process is compared withinitiation of the Izu–Bonin–Mariana Arc, whichformed by cross-cutting the older plate in theEocene (Uyeda & Ben-Abraham 1972) and subse-quently experienced back-arc spreading (Karig1971), and is also compared with subduction ofthese back-arc systems (Kyushu–Palau Ridge andShikoku Basin) beneath southwest Japan since theMiocene.
A serious problems in this model is to specify
the Cretaceous daughter arc, which might haveexisted ocean-ward of the Oku-Niikappu complex.Cretaceous–Paleogene island arc terranes are notuncommon in the western Pacific margins. Theyoccur in the Okhotsk to Kamchatka regions(Kuyul, Olyutorsky and Kronotsky terranes; Geistet al . 1994; Khudoley & Sokolov 1998; Saito et al .
1999; Silantyev et al . 2000), in Sakhalin (Schmidt
Peninsula and East Sakhalin Mountains; Raznitsin
1982; Rozhdestvensky 1986), in the Kurile OuterArc (Nemuro Belt; Kiminami 1983), in the DaitoRidge province in the western Philippine Sea(Shiki et al . 1977), and in Halmahera Island in
Indonesia (Hall et al . 1988), clustering around the
Okhotsk and Philippine Seas. As seen in Figure 11,the Kurile Outer Arc has collided with the centralHokkaido margin since the late Eocene–Oligocene(Fujiwara & Kanamatsu 1994; Arita et al . 2001),
after subduction of the young back-arc basin, andis thus a candidate for the daughter arc of the Oku-Niikappu remnant arc. Preliminary paleogeo-graphic sketches are displayed in Figure 12,although detailed stratigraphic and petrologic cor-relations between the arc terranes should be madein future.
Fig. 11 Ocean plate stratigraphy and reconstructed subduction history
for central to western Hokkaido. Hb, Horobetsugawa complex (Ueda et al.
2001); Nm, volcano-sedimentary sequence in the Nemuro Belt (Kiminami1983); Nz, Naizawa complex (Ueda et al. 2001); ONC, Oku-Niikappu
complex (Kiyokawa 1992 and the present study); Ru, Hidaka Belt accre-tionary complex, Rurochi Formation, by Kiminami et al. 1990, 1999; Sk,
ophiolites in East Sakhalin Mountains and Schmidt Peninsula, SakhalinIsland (Raznitsin 1982); To, Hidaka Belt accretionary complex, Tomu-raushi area, by Watanabe and Iwata 1987 and Miyashita and Katsushima1986. Ages of the Oshima Belt after the compilation by Kawamura et al.
(1997), and of the Sorachi Group after Kiminami et al. (1992).
Terrigenous & tuffaceous
sedimentary rocks
Chert
LimestoneAbyssal & back-arc basin
Inferred age of subducted oceanic crustIsland arcSeamount
Oceanic plateauAge of subduction and accretion
Late
EarlyMiddleLateEarlyEocene
PaleocenePg. Ng. Cret. Jurassic
LateTrias. Perm.C.EarlyMiddle
Late
Early
LateJurassic Paleogene Neog.
Late Early Middle Late Pc. Eocene Mioc. P. Oc.Cretaceous
In-situ
MORBAge of sedimentationMa
Old oceanic
crust
Old oceanic plate (Izanagi plate)Daughter arc
Old oceanic plate[Late Jurassic][mid-Cretaceous]
Izanagi plateBack arc basinONCOligoceneMiocene
Pacific plate subductionKurile arc collisionRemnant arc subductionYoung
oceanic crustOshimaIdonnappu
SorachiHidaka Nemuro
Hb
NzToRu
Nm
Sk50
100
150
200
250
300050 100 150 0
Kula or
Pacific
plateEurasian active
margin
Igneous and volcaniclastic rocks[Pre-Late Jurassic]
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Subducted remnant arc in Cretaceous Japan 595
UPPER PLATE RESPONSES TO REMNANT
ARC SUBDUCTION
The age of the oceanic crust differs discontinu-
ously across a remnant arc (Fig. 1A). This isbecause remnant arcs are formed by intermittentopening of back-arc basins (cf. West MarianaRidge), or by back-arc opening of island arcs thatformed by cross-cutting pre-existing old oceaniccrusts (cf. Kyushu–Palau Ridge). In any case, ifa remnant arc subducts (i.e. passes through atrench), the age of the subducting oceanic slabmust abruptly alter, leading to sudden changes inits physical, topographic and stratigraphic nature.
The condition of the subduction zone of central
Hokkaido has changed greatly since the timingof accretion of the Oku-Niikappu Remnant Arc(Fig. 13). From the Early Cretaceous to the LateCretaceous, subduction zone metamorphism in theKamuikotan zone transformed from a typical blue-schist facies series to a high-pressure–intermedi-ate facies series (Sakakibara & Ota 1994; Iwasakiet al . 1995; Ota 1999). Trench tectonics converted
from accretion of seamount volcanics with lesserturbidites, accompanied by tectonic erosion, toaccretion of voluminous trench-fill turbidites(Ueda et al . 2000, 2001). The volcanic front in the
upper plate shifted toward the back-arc side fromthe eastern margin of the Oshima Belt (Rebun–Kabato Sub-belt) of Hokkaido in the Early Creta-ceous to far-east Russia in the Late Cretaceous–Paleogene (Takahashi 1983). The discontinuous
younging of the subducted oceanic crust sincethe mid-Cretaceous (Fig. 11) can explain thesechanges through an increase of the thermal gradi-ent in the forearc, a decrease in the frequency of
seamount subduction, an increase in the slab buoy-ancy resulting in voluminous sediment supply , anda decrease in slab dip, which shifted the volcanicfront to the back-arc side.
A similar set of contrasting tectonic conditions
is also seen along the northern margin of the Phil-ippine Sea Plate (Fig. 1A). To the east of the cur-rently subducting Kyushu–Palau Ridge, Mioceneoceanic crust is subducting (Okino et al . 1994),
where the trench (Nankai Trough) is accompaniedby a clastic-dominant accretionary wedge. Thearc–trench gap exceeds 300 km owing to gentle dipsubduction (Nakamura et al . 1997). In contrast, toFig. 12 Preliminary paleogeographic model for the Cretaceous northwest Pacific. (A) The old Izanagi Plate was broken to form a new margi nal basin
plate and an intraoceanic arc, which subsequently suffered back-arc spreading in the Early Cretaceous. (B) The remnant arc subd ucted beneath the Eurasian
margin to become the Oku-Niikappu complex (ONC) in the mid-Cretaceous. (C) Back-arc spreading centers in the marginal basin pla te (or the descendent
Okhotsk Plate) subducted beneath the Eurasian margin in the Late-Cretaceous–early Paleogene, followed by the approach of the Ku rile–Kamchatka
Daughter Arc. Both the Kurile–Kamchatka and Izu–Bonin Arcs began to collide with Japan in accordance with the Japan Sea opening since the Oligocene–
Miocene (not shown).
C-P?K
Izanagi
Active arc
Remnant arcHalmahera &
proto-DaitoK
KK
JKamchatka
?Hawai
Kurile
& NE SakhalinKula
Izanagi[ Age of oceanic crust ]
Pg: Paleogene
K: CretaceousJ: Jurassic
T: Triassic
P: PermianC: CarboniferousDaito Ridge
provincePacificK
K
K
HalmaheraProto-
Izu-Boni
n-Ma
riana(Kyus
hu-P
alu Ridge
)Pg
Hawaii
ONC
KKKRemnant arc
subductionKula
Hawaii
JKPgPgK
A marginal basin plate
(separated from Izanagi)
KEurasiaP-T
Pg(A) 120–140 Ma
Idonnappu
Shimanto(ONC)
IdonnappuHidaka
Shimanto
Chichibu
?
???
Pacific(C) 40–50 Ma (B) 70–100 Ma
C-P
Nemuro-TokoroP-T
Fig. 13 Contrasting tectonic regimes in Cretaceous central Hokkaido
across the remnant arc subduction. (A) Early Cretaceous; (B) LateCretaceous.
Young oceanic
crust (Cretaceous)Basaltic-andesiticsubmarine arc
Andesitic-rhyolitic
continental arcSW Hokkaido
(Oshima Belt)Central Hokkaido
(Sorachi-Yezo Belt)Far-East
Russia
Remnant
ONC as suture zone
(accreted remnant arc)Seamount accretion
& tectonic erosion
Turbidite accretion
Gentl e-dip subductionSteep-dip subductionSorachi ophioliteJurassic
accretionary
complex
Pumpellyite-actinolite
facies metamorphism
(with calcite)Blueschist faciesmetamorphism(with aragonite)arc
Thermal gradient
increasedVolcanic front
retreatedAccreted materials changed
from seamounts to sedimentsA: Early Cretaceous
B: Late CretaceousOld oceanic crust
(Permo-Triassic: cold & dense)
Young oceanic crust
(Cretaceous:
warm & buoyant)
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596 H. Ueda and S. Miyashita
the west of the Kyushu–Palau Ridge, an Eocene or
older oceanic crust is subducting along the sedi-ment-starved Ryukyu Trench (Hilde & Lee 1984),and the arc–trench gap is as short as 200 kmreflecting its steeper-dip subduction.
Therefore, remnant-arc subduction plays an
important role in subduction zone evolution,because it can alter a subducting oceanic slab ofdiffering nature within a short period, promotingdrastic change in the upper plate tectonics.
CONCLUSIONS
1. The Oku-Niikappu complex originated from an
intraoceanic remnant arc, which formed byback-arc spreading of pre-existing island arc inthe earliest Cretaceous, and then accreted tothe mid-Cretaceous Eurasian continentalmargin.
2. Stratigraphic succession from arc volcanics to
pelagic sedimentary rocks, via arc-deriveddebrites in cases, is diagnostic for identifyingancient intraoceanic remnant arcs in orogenicbelts. Interbedded occurrences of pelagic sedi-mentary rocks and arc-derived debrites alsosuggest the same settings.
3. Accreted fragments of intraoceanic remnant
arc indicate subduction of marginal basinplates. A marginal basin plate separated froman old oceanic plate is proposed for the Creta-ceous northwest Pacific.
4. A remnant arc subduction abruptly alters the
age of the subducting oceanic lithosphere,resulting in drastic change in the upper platetectonics, such as accretion style, metamorphicconditions and volcanic front migration.
ACKNOWLEDGEMENTS
We thank the Hokkaido Electric Power Co., Inc.
for their facilities for our field survey , T . Nagahashifor supplying a spinel-bearing metavolcanic sam-ple, T . Katoh who provided information and litera-ture on Sakhalin Island, and K. Kizaki and T .Tomatsu for their help during our field survey .
Thanks are also due to the editors of this thematicissue (Y . Dilek and Y . Ogawa), and to K. Fujiokaand A. Ishiwatari for constructive reviews.H. Ueda was supported by the MEXT Grant-in-Aid for JSPS Fellows no. 01403900, and by the 21stCentury COE Program ‘Neo-Science of NaturalHistory’ (Hokkaido University).REFERENCES
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Earth Planets Space, 50, 153–163, 1998
153Copy right © The Society of Geomagnetism and Earth, Planetary and Space Sciences
(SGEPSS); The Seismological Society of Japan; The Volcanological Society of Japan;
The Geodetic Society of Japan; The Japanese Society for Planetary Sciences.1. Introduction
Akita-Yakeyama volcano is situated at the northwestern
edge of the Sengan Geothermal Area, northeast Japan,
where many geothermal features such as fumaroles and hotsprings are observed as well as many Quaternary volcanoes(Fig. 1). In late 1970s, an aeromagnetic survey was conductedin the area by (Ministry of International Trade and Industry(MITI), 1978); successive analyses of the magnetic anomalieswere conducted to make a regional magnetic model whichaccounts for the anomalies (MITI, 1978; Okuma and Suto,1987). Okuma and Suto (1987) calculated correlation co-efficients between the observed and synthetic magneticanomalies (Blakely and Grauch, 1983) and showed that therocks in the regions of positive correlation are normallymagnetized and those in the regions of negative correlationare reversely magnetized. They also conducted a spectralanalysis of the magnetic anomalies to estimate the depth tothe magnetic basement, indicating that the magnetic base-ment is shallower in the center of the area than in the edgesof the area. However, their analyses were focused primarilyon revealing the regional distribution of the direction of themagnetization of the volcanic rocks which compose themagnetic terrain and regional subsurface structure of thearea. Recently a new geothermal power plant began operationaround Akita-Yakeyama volcano as a result of intensivegeothermal explorations. Therefore, more detailed analysesand interpretations of the magnetic anomalies are necessaryfor revealing the detailed subsurface structure of the vol-cano and its surrounding areas. New techniques of theanalysis of the magnetic anomalies and accumulated othergeophysical data including paleomagnetic data make itpossible to conduct such a study.
2. Geology
Recently, detailed geologic maps have been compiled for
the Sengan Geothermal Area (Research group for the geo-logical map of Sengan Geothermal Area, 1985) and forAkita-Yakeyama volcano and its surrounding areas (Suto,1992), respectively. These maps show that the geothermalarea is covered widely by Quaternary volcanic rocks withoutcrops of Pleistocene ones (Fig. 2). Drilling showed thatthe surface volcanic rocks are underlain by Neogene volcanicand sedimentary rocks in the Sengan Geothermal Area. Theabsence of pre-Neogene rocks inside the geothermal areaand the presence of Paleozoic sedimentary rocks and Cre-taceous granitic rocks outside the area suggest the existenceof a large tectonic subsidence beneath the surface volcanicrocks.
According to an explanation of the geologic map of Akita-
Yakeyama volcano (Suto, 1992), the geology and tectonicsof the volcano are summarized as follows:
Akita-Yakeyama volcano is considered to be younger
than 1 Ma by K-Ar dating, a paleomagnetic study andvolcanostratigraphy. Most of the volcanic rocks from thevolcano are quartz and/or olivine-bearing pyroxene andesite,whereas small amounts of them are composed of dacite andrhyolite. A piston cylinder type depression is observed rightbeneath Akita-Yakeyama volcano, which is named as thePre-Yakeyama caldera, and is considered to have collapsedduring the eruption of the Tamagawa Welded Tuffs. An-Magnetic constraints on the subsurface structure of Akita-Yakeyama volcano,
northeast Japan
Shigeo Okuma
Geological Survey of Japan, Higashi 1-1-3, Tsukuba, Ibaraki 305-8567, Japan
(Received February 17, 1997; Revised September 24, 1997; Accepted November 21, 1997)
Magnetic analyses have been conducted in and around Akita-Yakeyama volcano at the northwestern edge of the
Sengan Geothermal Area, northeast Japan to reveal the regional and local subsurface structures of the area. First,
a magnetization intensity mapping method has been applied to analyze aeromagnetic anomalies of the area.Generally, magnetization highs and lows lie on volcanic rocks which are normally and reversely magnetized,respectively. Magnetization lows with small amplitudes are distributed on hydrothermally altered areas. These
results imply the usefulness of the method to estimate the young volcanic activities of Quaternary volcanic areas.
Detailed magnetic modeling reveals the subsurface structure of Akita-Yakeyama volcano itself. Rock magneticdata from volcanic rocks, both from the surface and cores in the geothermal exploration wells, have been employedfor the modeling. The resultant magnetic structure indicates the following: the surface volcanic rocks are underlainby granitic intrusions which have minimum thicknesses of about 2,000 m below the northern flank of volcano; inthe southern flank, the surface volcanic rocks are underlain widely by the Old-Tamagawa Welded Tuffs which are
reversely magnetized. These results show a good agreement with a geologic interpretation in and around the
volcano, especially with a hypothesis of the existence of buried calderas below the present volcano.
154 S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO
Fig. 1. Topographic map of the Sengan Geothermal Area. Contour interval is 100 m. The large and small boxes show the close-up and the detail study
area of Akita-Yakeyama volcano, respectively.
Fig. 2. Geological map of the Sengan Geothermal Area modified from Research group for the geologic map of Sengan Geothermal Area (1985).
Younger Volcanic Rocks 1: Matsukawa Andesite, Younger Volcanic Rocks 2: Otsuki-yama volcanics, Younger Volcanic Rocks 3: Hachim antai
volcanics and Kayo-dake volcanics, Younger Volcanic Rocks 4: Nanashigure-yama volcanics, Takakura-yama volcanics, Moriyoshi-yam a
volcanics, Akita-Yakeyama volcanics, Iwate-san volcanics and Akita-Komagatake volcanics.
S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO 155
other caldera is also recognized south of the volcano: the
Old-Tamagawa caldera. The caldera is thought to have
collapsed during the eruption of the Old-Tamagawa Welded
Tuffs in the Neogene, resulting two stage depressions filled
up with welded tuffs with the thicknesses exceeding 1,000
m.
3. Geologic Factors on Geothermal Explorations
Geothermal convection systems (White, 1967; Fig. 3) are
dominant in the Sengan Geothermal Area. Fracture systems
are expected to exist along fault systems and/or geologic
boundaries between the basement and overlaying layers.
Therefore, revealing concealed fault systems and structures
of the basement are important for geothermal exploration in
the area. A regional fault system, the Hanawa graben, lies
north and south between Akita-Yakeyama and Hachimantai
volcanoes. The Onuma geothermal power plant is located in
the center of the graben, while the Sumikawa geothermal
power plant lies at the western edge of the graben (Fig. 1;
Komazawa, 1987).
High-temperature geothermal fluids associated with
fractures cause hydrothermal alterations of the rocks which
compose volcanic terrains and are essential for power
generation in this area. Mapping of hydrothermally altered
areas on the surface has been conducted by geologists but
needs to be done beneath the surface.
Hot intrusive bodies are sometimes direct heat sources of
geothermal activities but are not common in the area except
Matsukawa area, where a geothermal power plant is under
Fig. 3. Generalized model of a geothermal convection system, modified from White (1967).operation by employing vapor dominated resources.
Important geologic factors on geothermal explorations in
the Sengan Geothermal Area are summarized as follows:
regional and local fault systems, structures of the basement,
hydrothermally altered areas, and hot intrusive rocks. In this
paper, I will discuss a mapping of hydrothermally altered
areas in the Sengan Geothermal Area by employing an
inversion of magnetic anomalies and a magnetic modeling
of the subsurface structure of Akita-Yakeyama volcano.
4. Magnetization Intensity Mapping
In conjunction with various geophysical surveys, an
aeromagnetic survey for geothermal exploration was con-
ducted in the Sengan Geothermal Area (MITI, 1978). Okuma
and Suto (1986) recompiled aeromagnetic anomaly maps
from the original anomaly data. The observed magnetic
anomalies (Fig. 4) are influenced strongly by the effect of
topographic relief. The volcanic rocks which compose the
terrain have their magnetic properties, strong enough to
cause intensive magnetic anomalies.
To remove the effects of magnetic terrain, I applied an
apparent magnetization intensity mapping method (Okuma
et al., 1994a; Nakatsuka, 1995) to the magnetic anomalies of
the Sengan Geothermal Area. These methods were useful
for revealing concealed old volcanic edifices on Izu-Oshima
Island (Okuma et al., 1994a) and the quadratic distribution
pattern of magnetization intensity along the Tanna fault,
presumably caused by movements of the strike-slip fault
(Nakatsuka, 1995).
156 S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO
Fig. 5. Magnetic model employed for an apparent magnetization intensity
mapping in this study. Jij: Magnetization of a prism at ij-th point, Tij:
Synthetic magnetic anomaly at ij-th point, ZijU: Top depth of a prism
atij-th point, ZijL: Bottom depth of a prism at ij-th point.Fig. 4. Total intensity aeromagnetic anomalies (IGRF residuals) in the Sengan Geothermal Area. Contour interval is 50 nT. The altitude of datum plane
is 1,800 m Above Sea Level. See also Fig. 1.
I assume a magnetic model composed of finite rectangular
prisms that corresponds to the volcanic terrain (Fig. 5). The
top depth of the model corresponds to the ground surface,
while the bottom depth corresponds to a flat surface, though
it can be assumed to be of any shape. In this case, the
synthetic magnetic anomaly Tx can be formulated by Eq. (1).
TW J x nxx i i
im
==() ( )
=∑
112 1 , , , K
where Wxi: Geometrical factor from the x-th point against
thei-th prism (Bhattacharyya, 1964); Ji: Magnetization of
i-th prism.
The observed magnetic anomaly Fx can be formulated in Eq.
(2).
WJ F x nxi i
im
x
=∑==() ( )
112 2 , , , . K
Ifm = n, it is straight forward to solve this equation because
it has a unique solution. Nevertheless, the solution sometimes
reveals an oscillation in an ill-conditioned case, probably
due to the accumulation of numerical error. Therefore, I
employ the conjugate gradient method (Nakatsuka, 1995)
instead of solving Eq. (2) directly. A flat surface of 5 km
below sea level was used for the bottom depth of the
magnetic model, taking account of a local shallow Curie
S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO 157
Fig. 6. Magnetization intensity map of the Sengan Geothermal Area with a topographic shading. Contour interval is 0.2 A/m. The areas bounded by
solid lines indicate hydrothermal altered areas. Solid and open circles locate volcanic rocks which are normally and reversely magnetized,
respectively. Paleomagnetic data (Suto, 1985, 1987, 1992; Suto and Mukoyama, 1987) which have stable magnetization intensities (≥1.0 A/m) with
magnetic pole latitudes ( ≥|±50°|) were plotted. See also Fig. 1.
Fig. 7. Magnetization intensity map of Akita-Yakeyama volcano with a topographic shading. The map area corresponds to the small box in Figs. 1
and 6. Contour interval is 0.2 A/m. Red broken lines indicate caldera rims of the Pre-Yakeyama caldera (north) and the Old-Tamagawa caldera
(south). See also Fig. 6.
158 S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO
depth (<8 km) (Okubo et al., 1989).
Characteristics of the distribution of the magnetization
intensity (Fig. 6) are summarized as follows:
Magnetization highs lie obviously on Quaternary volca-
noes such as Maemori-yama volcano, Iwate-san volcano,
Obuka-dake volcano, Takakura-yama volcano and Akita-
Komagatake volcano. This result clearly shows that each
edifice is roughly magnetized in a direction of the present
earth ’s magnetic field. On the other hand, obvious magne-
tization lows lie on Kurasawa-yama and Kayo-dake vol-
cano. Paleomagnetic dada indicate that normally magnetized
volcanic rocks are distributed in high magnetization areas,
whereas reversely magnetized rocks are in low magnetiza-
tion areas. This relationship is much clearer than that of
Okuma and Suto (1987).
Magnetization lows with small amplitudes, or rather to
say weak magnetization areas, correspond to hydrothermally
altered areas mapped by Research group for the geological
map of Sengan Geothermal Area (1985); weak magnetiza-
tions lie on and around the summit of Akita-Yakeyama
volcano, between Sumikawa and Onuma, in Matsukawa and
Kakkonda, between Takakura-yama and Akita-Komagatake
volcanoes, etc. It is implied that hydrothermal alterations
caused a loss of a large amount of magnetic minerals in
volcanic rocks. This kind of relationship was also observed
in Yellowstone National Park, USA (Okuma et al., 1994b).
Polygon No. Induced magnetization Remanent magnetization Top depth Bottom depth
Susceptibility
contrastInclination Declination Intensity
contrastInclination Declination ASL ASL
× 10–2 (SI) (degrees) (degrees) (A/m) (degrees) (degrees) (km) (km)
1 3.1 53.5 –7.5 0.0 0.3 –2.0
2 3.1 53.5 –7.5 0.0 0.3 –2.0
3 3.1 53.5 –7.5 0.0 –0.5 –2.0
4 0.0 2.0 –53.5 172.5 0.7 0.1
5 0.0 2.0 –53.5 172.5 0.1 –0.5Table 1. Parameters of horizontal polygons employed for analyzing magnetic anomalies of Akita-Yakeyama volcano.Fig. 8. E-W geologic cross section of the northern flank of Akita-Yakeyama volcano, inferred from drillings (after NEDO, 1988) .Recently, young granitic intrusions have been found in
Kakkonda area, one of weak magnetization areas, by drill-
ing. A temperature of the bodies exceeded 600 °C at the
bottom of a drill hole close to the Kakkonda geothermal
power plant. On the basis of the results, the apparent mag-
netization intensity method is useful to estimate the location
of hydrothermal areas in Quaternary volcanic areas.
5. Forward Modeling for Magnetic Structure of
Akita-Yakeyama Volcano
Next, I will focus especially on the magnetic anomaly
(Fig. 4) and the resultant magnetization intensity (Fig. 7) in
and around Akita-Yakeyama volcano. The distribution pat-
tern of magnetization intensity of the volcano is different
totally from that of typical young volcanoes in the Sengan
Geothermal Area. Usually a magnetization high is distrib-
uted on the edifice of each volcano. Instead, a magnetization
low lies right on the southern flank of the volcano, surrounded
by magnetic highs.
In conjunction with rock magnetic data from surface
volcanic rocks and core samples from geothermal explora-
tion wells, I will try to construct a magnetic model which
accounts for the magnetic anomalies. Drilling showed the
existence of Tertiary granitic intrusions at an altitude of sea
level below the northern flank of the volcano (NEDO, 1988;
Fig. 8). Rock magnetic measurements indicate that the
S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO 159
granitic rocks have the magnetic susceptibility (4 π× 2.5 ×
10–3 SI) strong enough to cause intensive magnetic anoma-
lies, though the natural remanent magnetization (NRM) of
the rocks is negligibly small ( ≤2× 10–1 A/m) (NEDO, 1986a;
Fig. 9). Drilling also showed that the surface volcanic rocks
on the southern flank of the volcano are underlain by theOld-Tamagawa Welded Tuffs (NEDO, 1986b). According
to paleomagnetic measurements (Suto, 1985, 1987, 1992;
Suto and Mukoyama, 1987), most of the Old-Tamagawa
Welded Tuffs are reversely magnetized as well as the
Tamagawa Welded Tuffs to the south.
Magnetic anomalies which are caused only by the volca-Fig. 9. Columnar section of geothermal exploration wells, SN-5, SN-7D, SN-6K and SN-2, in the northern and southern flanks of Akita-Yakeyama
volcano. Vectors show inclinations of NRM of rock specimens from each unit, except for that of SN-7D (paleomagnetic inclination). A.P.: Pre-
Yakeyama Andesite Pyroclastic Rocks, M.F.: Mataguchigoya Formation (Pyroxene dacite tuff) (after Suto, 1992). See also Figs. 2 and 8.
160 S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO
Fig. 11. Magnetization intensity map of Akita-Yakeyama volcano and its
vicinity. Contour interval is 20 × 10–2 A/m. Solid and broken lines
show positive and negative values, respectively. See also Fig. 10.
nic edifice (Fig. 10) were extracted from the observed
anomalies (Fig. 4). The magnetic anomalies (Fig. 12) were
calculated from the magnetization intensity data within an
area bounded by a thick solid line (Fig. 11). Next, synthetic
anomalies (Fig. 13) caused by a terrain model (Fig. 10) with
a constant thickness of 500 m and a top corresponding to the
ground surface were calculated by taking into account of the
existence of hydrothermally altered areas: the altered areaswere assumed to be non-magnetic. Then, residual anomalies
(Fig. 14) were calculated by subtracting synthetic terrain
anomalies (Fig. 13) from the extracted observed anomalies
(Fig. 12). Finally, optimal synthetic anomalies (Fig. 15)
were calculated to fit the extracted observed anomalies (Fig.
12) by employing five horizontal polygons (Talwani, 1965).Fig. 12. Extracted total intensity aeromagnetic anomalies (IGRF residu-
als) of Akita-Yakeyama volcano. The anomalies were calculated from
prism models within a thick solid line in Fig. 10, which have own
magnetization intensities in Fig. 11 and have a flat bottom at an altitude
of 5 km below sea level. Contour interval is 10 nT. See also Fig. 11.Fig. 10. Topographic map of Akita-Yakeyama volcano and its vicinity.
The map area corresponds to the large box in Figs. 1 and 6. The
patterned area bounded by thick solid lines indicates a terrain model
which has a constant thickness of 500 m and its top is the ground
surface, excluding hydrothermal altered areas which correspond to non
magnetic areas. Contour interval is 100 m. See also Fig. 6.
Fig. 13. Synthetic total intensity magnetic anomalies calculated from a
terrain model with a boundary denoted by a thick solid line in Fig. 10.
Magnetization intensity of the terrain model is assumed to be uni-
formly 2.0 A/m except for the local hydrothermally altered areas which
are non-magnetic. Contour interval is 10 nT. See also Fig. 11.
S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO 161
The goodness-of-fit ratio ( r) (Blakely, 1995) was employed
to indicate a goodness of fittings between two data sets.
rF Rx
xn
x
xn
= ()
==∑∑
113
where, Fx: Observed magnetic anomaly at the point of x;Rx:
Residual anomaly between the observed and syntheticanomaly at the point of x.
In this equation, the larger ratio ( r) shows the better fitting.
I employed the magnetic parameters (Table 1), determined
from rock magnetic measurements (NEDO, 1986a, b; Okuma
and Suto, 1987; Okuma, in preparation). The resultant
magnetic model is composed of five polygons (Table 1;
Figs. 15, 16(a), and 16(b)): three of them correspond to
Tertiary granitic intrusions beneath the northern flank, while
the rest corresponds to the reversely magnetized volcanic
rocks such as the Old-Tamagawa Welded Tuffs beneath the
southern flank.
6. Discussions
I discuss the results of the magnetic modeling for the
subsurface structure of Akita-Yakeyama volcano with other
information.
Tertiary granitic intrusions were found at the bottoms of
geothermal exploration wells, SN-5 and SN-7D, beneath the
northern flank of Akita-Yakeyama volcano but the bottom
depths remain unknown. In this area, the corresponding
granitic intrusions are believed to have relation to active
geothermal activities; a maximum lost circulation was ob-
served at the boundary between the intrusions and the
overlaying layers, where a major geothermal reservoir is
expected to exist (Nish et al ., 1989). Therefore, it is im-
portant to get information about the extent of the intrusions.
Polygon 1, 2 and 3 lie just inside the northern rim of the Pre-
Yakeyama caldera (Figs. 7 and 15), suggesting that the
caldera collapsed along the faults between the intrusions and
the surrounding layers.
Geothermal exploration wells, SN-5 and SN-7D, where
high temperatures more than 200 °C were observed, are
located just above horizontal boundaries of polygon 2 and 3,
respectively (Fig. 15), implying the existence of concealed
faults and fractures where hydrothermal fluids can flow
easily. Resistivity models were obtained by a two-dimen-
sional inversion of MT, EMAP (Electromagnetic Array
Profiling) and CSAMT data in northern flank of Akita-
Yakeyama volcano (Uchida and Mitsuhata, 1995). A rather
high resistivity in an order of 100 Ω·m, which corresponds
to the granitic intrusions, was analyzed beneath the area and
extends at least 2 km in a north-south direction, supporting
the results of the magnetic modeling. As for geothermal
explorations, boundaries between the intrusions and the
surrounding layers should be a target for drilling, because
fracture zones are highly expected.
The minimum bottom depths of polygon 1, 2 and 3
beneath the northern flank are about 2,000 m below sea
level, if I fix the top depths at and around sea level on the
basis of geologic information obtained from drilling (Figs.
15 and 16; Table 1). A temperature over 300 °C was observed
at the bottom of a geothermal exploration well, SN-7D,
(NEDO, 1988), implying a demagnetization of the granitic
rocks. If so, the bottom depth of these models might corre-
spond to the local Curie depth in this area.
Geothermal activities, such as hot springs and hydrother-
mal alterations, are less active in the southern flank than in
the northern flank of the volcano. This makes the southern
flank become less attractive from a viewpoint of geothermal
explorations. However, the subsurface structure of theFig. 14. Residual total intensity magnetic anomalies between the ob-
served (Fig. 12) and synthetic magnetic anomalies (Fig. 13). Contour
interval is 10 nT. See also Fig. 11.
Fig. 15. Synthetic total intensity magnetic anomalies which best fit the
residual anomalies (Fig. 14). Contour interval is 10 nT. Goodness-of-
fit ratio is 1.8. Numbers 1 –5 denote horizontal polygons. Circles show
the locations of geothermal exploration wells. See also Fig. 11.
162 S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO
(b)(a)
Fig. 16. (a) N-S cross section of synthetic total intensity magnetic anomalies (Fig. 15) with the magnetic model. See also Fig. 15. (b) E-W cross section
of synthetic total intensity magnetic anomalies (Fig. 15) with the magnetic model. See also Fig. 15.
southern flank, as well as that of northern flank, is important
to better understand how the present edifice has formed.
Two polygons 4 and 5, which are reversely magnetized,
are necessary to account for the magnetic anomalies of the
area (Figs. 15 and 16(a)). Polygon 5 occupies the almost
southern half of the subsurface of the volcano, if I fix the top
and bottom depths on the basis of geologic information
obtained from drilling. This might indicate an evidence of
the northern extension of the Old-Tamagawa caldera be-
neath the southern flank of the volcano, though the exten-
sion inferred from drilling information of SN-6K (Suto,
1992; Fig. 2) is much smaller.
I have to emphasize the existence of polygon 4, because
without this it is very difficult to explain the corresponding
anomalies. This result implies the existence of a concealed
old volcano associated with reversely magnetized rocks
such as the Old-Tamagawa Welded Tuffs and/or the Pre-
Yakeyama Andesite Pyroclastic Rocks, which were found
only in drill holes such as SN-2 and SN-6K in the vicinity of
the volcano (Suto, 1992; Fig. 9). Actually, a local topo-
graphic high, Kuroishimori, exists on the southern flank
(Fig. 7), though the area is covered widely by YoungerVolcanic Rocks (Fig. 2). The summit of the main edifice of
the volcano is located just above the northern edge of
polygon 5, suggesting relationship with a concealed caldera
boundary.
7. Conclusions
In order to better understand a regional magnetic struc-
ture, I applied an apparent magnetization intensity mapping
to the magnetic anomalies in the Sengan Geothermal Area.
The resultant magnetization intensity map shows that mag-
netization highs and lows correspond to the distribution
areas of normally and reversely magnetized volcanic rocks,
respectively. Magnetization lows with small amplitudes
correspond to hydrothermally altered areas, suggesting a
loss of magnetic minerals which compose volcanic terrain.
These results lead us to a conclusion that the magnetization
intensity mapping is useful for plotting the distribution of
normally or reversely magnetized rocks and finding hydro-
thermally altered areas on Quaternary volcanic terrains.
A magnetic modeling, with known structural and mag-
netic parameters, was conducted to reveal a detailed subsur-
face structure of Akita-Yakeyama volcano. The magnetic
S. OKUMA: MAGNETIC CONSTRAINTS ON THE SUBSURFACE STRUCTURE OF AKITA-YAKEYAMA VOLCANO 163
model is composed of five polygons: three of them corre-
spond to granitic intrusions beneath the northern flank,while the rest of them corresponds mainly to the buried Old-Tamagawa Welded Tuffs beneath the southern flank. Thenorthern polygons show the depth extent of the graniticintrusions or the local Curie depth. The southern polygonsindicate a subsurface convex structure, implying the exist-ence of a concealed old volcano associated with the Old-Tamagawa Welded Tuffs and/or the Pre-Yakeyama AndesitePyroclastic Rocks.
Acknowledgments. The author wishes to thank Shigeru Suto,
Geological Survey of Japan, for his helpful criticism of the
manuscript. The author also would like to express his appreciation
to two reviewers, Carol Finn and Hidefumi Watanabe, for their
help to improve the manuscript.
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|
Okuma (1998) magnetic constraints on subsurface NE japan.txt
|
Petrological diversity and origin of ophiolites in Japan and Far East
Russia with emphasis on depleted harzburgite
AKIRA ISHIWATARI 1, SERGEI D. SOKOLOV 2 & SERGEI V. VYSOTSKIY 3
1Department of Earth Sciences, Faculty of Science, Kanazawa University, Kanazawa 920-
1192, Japan (e-mail: geoishw@kenroku, kanazawa-u.ac.jp)
2Geological Institute, Russian Academy of Sciences, Pyzhevsky 7, Moscow 109017, Russia
3Far East Geological Institute, Russian Academy of Sciences, Prospect 100 letiya 159,
Vladivostok 690022, Russia
Abstract: Ophiolites are divided into lherzolite-type (L-type) and harzburgite-type (H-type) by
the lithology of their mantle peridotites. Rare depleted harzburgite-type (DH-type) is
distinguished from the normal H-type by the more refractory nature of its mantle peridotite and
the occurrence of orthopyroxene-type cumulate rocks including iron-rich harzburgite and
orthopyroxenite. The Shelting (Sakhalin) and Krasnaya (Koryak Mountains) ophiolites in Far
East Russia, which have both depleted harzburgite and orthopyroxene-type cumulate rocks,
belong to this newly defined DH-type. The ophiolites in SW Japan-Primorye, NE Japan-
Sakhalin, and the Koryak Mountains in the northwestern Pacific margin have diverse ophiolite
types ranging from L- to DH-types. The wide petrological diversity, the common occurrence of
DH-type, and the presence of thick crustal sections in these ophiolites suggest regionally
inhomogeneous, commonly very high degrees of mantle melting over subduction zones, as in
the modern Mariana forearc environment. The ophiolites of Japan and Far East Russia range in
age from Early Palaeozoic to Cenozoic and are tectonically underlain by younger blueschists
and accretionary complexes. The spatial association of these ophiolites with blueschists is
analogous to the ophiolite-blueschist assemblages recovered from the Mariana forearc. This
association might have formed in a period of non-accretion at an oceanic subduction zone that
was followed by voluminous accretion of sediments, facilitating subsequent uplift of the
ophiolites and blueschists.
The concept of ophiolites as assemblages of mafic
and ultramafic igneous rocks formed in deep
ocean was first suggested by Steinmann (1927),
and was further developed into the current model
of obducted fossil oceanic crust-mantle by
Moores (1969), Coleman (1971) and Moores &
Vine (1971). This model helped to explain spread-
ing processes at constructive plate boundaries and
tectonic emplacement processes at destructive
plate boundaries as the theory of plate tectonics
was formulated. Ophiolites are currently used to
help identify suture zones and ancient continental
collision zones where they rest tectonically on
older continental crust. Coleman (1986) called
these ophiolites 'Tethyan-type' and distinguished
them from those in Circum-Pacific belts, which
were labelled 'Cordilleran-type'. In general, Cor-
dilleran-type ophiolites are incomplete, dismem-
bered and metamorphosed; however, they are none
the less useful for unravelling global tectonic
problems. The first purpose of this paper is to
show that the Cordilleran-type ophiolites in the
northwestern Pacific margin are emplaced over
From: DILEK, Y. & ROBINSON R T. (eds) 2003. Ophiolites
Special Publications, 218, 597-617. 0305-8719/03/$15 9 young continental crust (or an accretionary com-
plex younger than the ophiolite), are associated
with blueschists, and represent remnants of fossil
oceanic subduction zones rather than continental
collision zones.
Boudier & Nicolas (1985) classified ophiolites
into lherzolite- and harzburgite-types (designated
LOT and HOT, respectively, by Nicolas (1989)).
They postulated that HOT form at mid-ocean
ridges with relatively fast spreading rates. How-
ever, recent studies on contrasting structural fea-
tures of the Cyprus and Oman ophiolites suggest
that these two bodies, both HOT, could be corre-
lated with slow- and fast-spreading ridges, respec-
tively (Dilek et al. 1998). Petrological and
geochemical studies have revealed that at least
some Tethyan ophiolites were generated from
suprasubduction zone (SSZ) magmas rather than
mid-ocean ridge basalt (MORB) (e.g. Cyprus:
Miyashiro 1973; Robinson et al. 1983; Hrbert &
Laurent 1990; Oman: Alabaster et al. 1982;
Umino et al. 1990; Lachize et al. 1996; Ahmed &
Arai 2002; Bay of Islands: Edwards 1995; Suhr &
in Earth History. Geological Society, London,
The Geological Society of London 2003.
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598 A. ISHIWATARI ETAL
Edwards 2000). Pearce et al. (1984) classified
most HOT as SSZ-type and considered all LOT as
MORB-type. Ishiwatari (1985a) divided harzbur-
gite-type ophiolites into ordinary harzburgite (H)
and depleted harzburgite (DH) categories (Fig. 1),
and pointed out that DH-type ophiolites include
olivine (Ol)-orthopyroxene (Opx) cumulates. In
contrast, L- (lherzolite) and H-type ophiolites have
olivine-plagioclase (PI) and olivine-clinopyroxene
(Cpx) cumulates, respectively. Although the crys-
tallization order of minerals may change with
pressure, all ophiolitic cumulates are believed to
have crystallized at low, plagioclase-stable pres-
sures, and compositional variations of cumulus
minerals suggest that the crystallization sequence was not pressure dependent (Ishiwatari 1985a).
The associated volcanic rocks also typically show
similar chemical variations.
The residual peridotite of ophiolitic mantle
sections is generally homogeneous over a scale of
kilometres, if we exclude samples from the Moho
transition zone and from dykes or veins of dunite,
pyroxenite or gabbro. The mineral chemistry of
such residual peridotite is the most reliable mea-
sure of the degree of melting of the mantle
section. Variations in the Cr-number (Cr/(A1 §
Cr)) of spinel in the residual peridotite correspond
closely to variations in lithology and mineral
chemistry. The Cr-number is 0.3-0.5 (or less) in
L-type, 0.5-0.7 in H-type, and 0.7-0.9 in DH-
Ophiolite Type L type H type DH type
Volcanic Rocks
Cumulate Type
Cpx "502 in
mafic rocks
Crystallization
sequence of
maflc and
(Seismic Moho)
ultramaflc
cumulates N-, T-, E-MORB,
alkali basalt
Plagioclase type
0.8 wt.%
OI PI Cpx Opx N-, T-, E-MORB,
Island arc basalt
and andesite
Clinopyroxene type
0.4 wt.%
OI PI Cpx Opx
III Island arc basalt
and andesite,
boninite
Orthopyroxene type
0.1 wt.%
OI PI Cpx Opx
III
Residual mantle
peridotite
Bulk AI~3+CaO
Spinel Cr#
Opx AI203
Olivine Fo
World classic
examples Lherzolite and Cpx-
rich harzburgite
3-5wt.%
0.3 - 0.5
2-4wt.%
90 or lower
Liguria, Alps, Trinity
Bay of Islands* Cpx-bearing
harzburgite
1 - 2 wt.%
0.5 - 0.7
1 - 2 wt.%
90 - 91
Semail (Oman),
Troodos, Vourinos Cpx-free (depleted)
harzburgite
0- lwt.%
0.7 - 0.9
0- lwt.%
92 or higher
Papua, Adamsfield,
Khan Taishir
SW Japan-
Primodye
NE Japan-
Sakhalin
Koryak Mts.
Taigonos P.
Koryak Mts.
Mainits Zone Oeyama
Poroshiri, Nukabira
Povorotny
Elistratova South
Tamvatney
Yagel Yakuno
Khanka
Iwanai
Miyamori
Elistratova North
Chirinai? (Omi chromitite?)
Horokanai, Taka-
domari, Shelting
Nablyudeny
Krasnaya
Srednaya
Fig. 1. Petrological types of ophiolites, after Ishiwatari (1991), and examples of each type in the northwestem Pacific
margin. See Ishiwatari (1985a, 1991) for references. *Lewis Hills massif of the Bay of Islands ophiolite shows a
more depleted nature (Edwards 1995).
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 599
type ophiolites (Fig. 1), although the associated
dunite and chromitite may have more Cr-rich
spinel. An increase in spinel Cr-number with
degree of melting was demonstrated experimen-
tally by Jaques & Green (1980) and was related to
peridotite compositions by Dick & Bullen (1984).
Although the melting process may be complicated
by hydrous remelting of already depleted mantle
(e.g. Bloomer & Hawkins 1987; Sobolev &
Danyushevsky 1994) and later reaction with per-
colating melts (e.g. Cannat et al. 1990; Arai et al.
1996), the spinel Cr-number of regionally homo-
geneous residual peridotite probably reflects the
degree of melting of the primary fertile mantle.
Therefore, we consider the spinel Cr-number of
the residual peridotite as a significant parameter
for classifying ophiolites in this paper.
The second purpose of this paper is to report
new examples of DH-type ophiolites in the north-
western Pacific region, where L- and H-types are
also abundant, and to consider the geological
significance of the extreme diversity of the ophio-
litic mantle in this region.
The geological and petrological data on the
Russian ophiolites presented in this paper are
mostly based on co-operative field studies in
Primorye (1993, 1998, 1999 and 2002) and the
Taigonos Peninsula (1995 and 1997) as well as on a 1990 field workshop in the Mainits Zone,
Koryak Mountains (Bryan 1991). Some of the
important chemical data and geological maps
reproduced here are from publications in Japan
and Russia.
Ophiolites in the northwestern Pacific
margin
The northwestern Pacific margin, which extends
from Japan to Russia, contains numerous ophioli-
tic bodies commonly associated with blueschists
(Fig. 2). These ophiolites are highly variable in
age, lithology and chemical composition, even in
an apparently continuous, 'single' belt. The ages
of the ophiolites and associated blueschists are
summarized in Table 1.
Ophiolite belts in SW Japan and Primorye
The Palaeozoic-Mesozoic accretionary complexes
and associated ophiolites and blueschists of south-
western Japan may have been linked with the
Sikhote-Alin Mountains of Russian Primorye
before the Miocene opening of the Japan Sea
(Ishiwatari & Tsujimori 2003). Jurassic accretion-
ary complexes in Japan (Tamba, Mino and Ashio
Major ophiolite belts in the
northwestern Pacific margins
1000 km N50"
CHINA"!!~
N40 ~ ):)(:~
Oa=e-=u= i!!ii metamorphic belt
S ~,....,-. o- \
." KOREA Vla~
s ...i:.:."..~;~.i:-; \
S ~ IAnA
,=~g o o
E. Taiwan Ishigaki Yakuno / \
~ Island ISWJapan-PrimoryelMikabu
20 = ophiolite belts I N70~..,....,~ E140"
E120" . .......:..:::.::::::::i::!ii!ii!.ii!ii!i:.i~:~::.:::...
N60"~!i!iiiiii!iiii..:ii:i:ii!!iiiiiPenzhinaZone El60~ E180"
\ (Pekulney, Ust-Belaya! J~_ RUSSIA:: :::::::::::::::::::::::::::::::::::::::: and Kuyu! from N to S) Koryak Mountains
Peninsula ( Taigonos ..... '""~"~"";Ei~"iralovaii!ii!~ ophiolite belts
Kengevayam' Nablyudeny and from N to S) \i:il
(Tamvatney, Yagel.
~ Krasnaya and
Srednaya from Susunat ~ ~/ o o ~ NWto SE)
amuikotan (Horokanai, Takadomari,
wanai and Nukabira from N to S)
(Poroshiri and Horoman)
I NE Japan-Sakhalin I Pacific Ocean
ophiolite belts I
Fig. 2. Location of major ophiolite belts and ophiolite complexes in the northwestern Pacific margin. A possible
western extension of the Palaeozoic ophiolite-blueschist belt of the SW Japan-Primorye area is also shown (see
Ishiwatari & Tsujimori (2003) for discussion).
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600 A. ISHIWATARI ETAL
#
4,
E
P~
m m o
~z ~o ~ .~ I
m o
o o
o ~m
'= ~.~
m ,<
"~ 0
r ~n
9 ,.,~ ~
< o
r <~o
m o o
~ i
o
g g o~
o
o
e~
dP zones) are believed to correlate with the Samar-
ka-Nadanhada terranes in Primorye and north-
eastern China (Kojima 1989).
Palaeozoic ophiolites in southwestern Japan
include the Ordovician Oeyama ophiolite and the
Permian Yakuno ophiolite. The hornblende K-Ar
age of the Oeyama ophiolite is 450 Ma (Nishi-
mura 1998), whereas that of the Yakuno ophiolite
is 280 Ma (Shibata et al. 1977). The latter age is
in agreement with zircon U-Pb dates from the
same body (Herzig et al. 1997).
The Oeyama ophiolite is composed of lherzoli-
tic mantle peridotite (Cr-number 0.3) in the east-
ern Chugoku area (Kurokawa 1985) and of
clinopyroxene-bearing harzburgite (Cr-number
0.5) in the western Chugoku area. Podiform
chromitites (Cr-number 0.5) encased in dunite
bodies are abundant in the western area (Arai
1980; Matsumoto & Arai 1997) (Fig. 3a). In both
areas the mantle peridotites commonly contain
vermicular spinel-pyroxene aggregates. The man-
tle peridotites are thrust over the 320-280 Ma
Renge blueschist (Tsujimori & Itaya 1999), the
Permian Akiyoshi accretionary complex, and the
Permian Yakuno ophiolite. Small mantle peridotite
bodies of analogous nature are also present in the
Hida marginal belt to the east, where they are
associated with the Renge blueschist and eclogite
in the Omi area (Tsujimori et al. 2000). Orbicular
chromitite with high Cr-number (0.76-0.85) is
present in peridotites in the Omi area (Yamane et
al. 1988). The high Cr-number suggests formation
from highly refractory melts, although the degree
of melting of the mantle peridotite is generally
low (Fig. 3a). This ophiolite includes minor
gabbroic rocks and ultramafic cumulates (Kuro-
kawa 1985), and granulite-facies spinel metagab-
bro recrystallized to kyanite-bearing epidote
amphibolite has been reported from the Oeyama
massif (Ysujimori & Ishiwatari 2002).
The Permian Yakuno ophiolite has a complete
succession with MORB-type basalt, clinopyrox-
ene-type cumulates, and mantle peridotite consist-
ing of relatively depleted harzburgite with a
variable composition (Cr-munber 0.4-0.8) (Ishi-
watari 1985a). Although regional variations in the
mantle peridotite cannot be confimaed because of
limited exposure, the gabbroic and basaltic rocks
range from T-MORB in the eastern part to island
arc tholeiite in the western part (Ishiwatari et al.
1990) 9 In gabbros the Mg-number of clinopyrox-
ene varies from 0,85 in the eastern part to 0.7 in
the westem part (Fig. 4a) although plagioclase has
a relatively uniform composition (An80) through-
out. Another peculiar feature of the Yakuno
ophiolite is the occurrence of granulite-facies
metacumulates (e.g. spinel-two-pyroxene meta-
gabbro) at the Moho level, which suggests unu-
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 601
(a) SW Japan-Primorye (b) NE Japan-Sakhalin
Khanka ophiolite, Primorye
~ = in conglomerate Opx
r~ in serpentinite 6.0- AI203
i i i i 110 0.8 0.6 0.4 0,2 0'.0 _wt'% ~_./H~176
Spinel Cr# Lh ~*~ .~/
4.0- Opx 4 0 Nukabira AI203
wt.% @ Oeyama E " /,~--%2t'L,',,~L. Nukabira Hz/~/
- amaW/3-V~ - . / ,~ ~ Mqamori Oey ~wana~ ~ :, .// 2.0- ,,.u- domari.~ Miyamori
.~'/Yakuno Horoman....~ ~~ Cr-.sp
BDH - ~ 2,~/~.d~_.-- Horokanai Hz
0.0 , ~_ Omi, Cr, , n n ..~_.~-~---Cape Shelting
i i i i i i v .,,./ i i i i i i i i i i
.0 0.8 0.6 0.4 0.2 0.0 1.0 ~0.8 0.6 0.4 0.2 0.0
Spinel Cr# Horokanai Du Spinel Cr#
6.0-
4,0-
2.0-
0.0
1.0 (c) Taigonos Peninsula (d) Mainits Zone, Koryak Mountains
Opx 6.0- Opx
AI203 AI203 Tamvatney Lh
wt.% Povorotny Lh _ wt.% ~'~
Tamvatney Hz
.~ Elistratova S 4.0- Yagelge~ melan
E listratova~/
//,,,.~/
Nablyu-/~ 2.0- /T.~/ "-_ Povorotny Hz deny
=(-~ - "= Mt. Krasnaya
,. Mt. Srednaya Du
08 06 04 o'2oo o o , ' ' ' 1,0 018'016'014'012 0',0
Spinel Cr# Spinel Cr#
Fig. 3. Compilation of spinel Cr-number and A1203 content of coexisting orthopyroxene in residual mantle peridotite
of ophiolites in (a) SW Japan-Primorye area, (b) NE Japan-Sakhalin area, (c) Taigonos Peninsula and (d) Mainits
Zone, Koryak Mountains. References: (a) Arai (1980), Ishiwatari (1985a, 1985b), Kurokawa (1985), Yamane et al.
(1988), Matsumoto & Arai (1997), Shcheka et al. (2001); (b) Ishizuka (1985, 1987), Ozawa (1988), Takahashi
(1991), Tamura et al. (1999), Vysotskiy et al. (2000); (c) Saito et al. (1999), Bazylev et al. (2001); (d)Dmitrenko et
al. (1990) and this study. HMLS, main harzburgite-therzolite suite; BDH, banded durite-harzburgite.
sually thick (15-30 km) oceanic crust (Ishiwatari
1985b).
In Primorye, the Sergeevka, Kalinovka and
Bikin ophiolitic complexes lie along the western
margin of the Mesozoic accretionary complexes
near the boundary of the Khanka crystalline
massif (Fig. 2). These ophiolites lack mantle
peridotite and are composed solely of volcanic
rocks, gabbros and mafic-ultramafic cumulates. On the other hand, the Cambrian Khanka ophiolite
rests on a Cambrian limestone in a small rift zone
(Spassk zone) in the Khanka massif. Middle
Cambrian conglomerate covering the body con-
tains abundant chromian spinel grains derived
from the ophiolite (Shcheka et al. 2001). This
ophiolite contains serpentinized harzburgite with
relatively chromian, unusually Mn-rich spinel (Cr-
number 0.6-0.7) (Shcheka et al. 2001); it is
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602 A. ISHIWATARI ETAL
0.4 I I I I I I I I [
(a) Yakuno ophiolite, SW Japan
Western Area
0.6- (Kamigori, Ohara) 9 Island-arc basalt
Eastern Area
(Oi-Maizuru) o
o0.8
MORB
1.0 I I I I I I L I I
0 20 40 60 80 100
Plagioclase An%
0.4 I ] I I I I I I I
(b) Taigonos Peninsula, NE Russia
=~0~06 9 Elistratovadiabasegabbro ophiolite ~lsland'arc basalt
Kengeveem gabbro o
~0.8 o
MORB
1.0 I I I I I I I I I
0 20 40 60 80 100
Plagioclase An%
Fig. 4. Relationship between anorthite (An) content of
plagioclase and Mg-number (= Mg/(Fe + Mg)) of
coexisting clinopyroxene in the Yakuno ophiolite in SW
Japan (a) and among the Elistratova ophiolite and
Kengevayam gabbro body, Taigonos Peninsula, NE
Russia (b). Fields for island arc basalts (IAB) and mid-
ocean ridge basalts (MORB) are from Ishiwatari et al.
(1990). Beard (1986) and Hrbert & Laurent (1990)
proposed a similar discrimination between the two
suites.
considered to be an H-type ophiolite transitional
to DH-type (Fig. 3a). The residual peridotite
spinels of the Khanka and Oeyama ophiolites
suggest that the early Palaeozoic mantle of the
SW Japan-Primorye belt was inhomogeneous and
was locally relatively depleted.
Dobretsov et al. (1994) correlated the Oeyama
ophiolite with the Kalinovka ophiolite; however,
age and structural data indicate that it is part of
the Sergeevka terrane, a huge metagabbro body
covered by Devonian and Permian sedimentary
rocks. The Sergeevka terrane was later thrust onto
a Jurassic accretionary complex (Samarka terrane)
containing 250 Ma blueschist. The Sergeevka ter-
rane, however, does not contain any mantle peri-
dotite, which is abundant in the Oeyama ophiolite.
The Bikin ophiolite, which consists of granulite-
facies two-pyroxene metagabbro (Vysotskiy 1994)
may be the Russian counterpart of the Permian
Yakuno ophiolite. Recent geochronological data suggest that the Yakuno ophiolite may also corre-
late with the Kalinovka ophiolite, which includes
spinel-bearing troctolite and garnet-beating meta-
gabbro (Ishiwatari & Tsujimori 2003).
Ophiolite belts" in NE Japan and Sakhalin
The early Palaeozoic Miyamori (and Hayachine)
ophiolite is present in the Kitakami Mountains,
NE Honshu (Ozawa 1994). This ophiolite and its
overlying Silurian to Jurassic sedimentary rocks
were thrust over the Jurassic accretionary complex
of the North Kitakami-SW Hokkaido zone (Taza-
wa 1988). The ophiolite consists mainly of moder-
ately depleted harzburgite (spinel Cr-number 0.4-
0.75) and clinopyroxene-type cumulates. The per-
vasive occurrence of hornblende (and some Yi-
poor phlogopite) in the mantle peridotite indicates
that it was derived from a hydrous mantle wedge
above a subduction zone (Ozawa 1988). The
harzburgite contains layered harzburgite-wehrlite
zones, which formed by melt-mantle interaction
(Ozawa 1994). The Miyamori ophiolite is asso-
ciated with the Motai blueschist of Late Palaeo-
zoic age (Maekawa 1988).
In Central Hokkaido, ophiolitic rocks occur in
several zones. In the Sorachi-Yezo belt, the
Jurassic Horokanai ophiolite is thrust over the
Cretaceous Kamuikotan metamorphic belt, and
dismembered equivalents of this ophiolite are
distributed throughout the belt. The Horokanai
ophiolite has a well-preserved succession, which
is composed, from bottom to top, of harzburgite,
orthopyroxenite-dunite cumulate rocks, metagab-
bros, amphibolite, MORB-type pillow lavas and
an Upper Jurassic radiolarian chert. The mantle
harzburgite is strongly depleted (spinel Cr-number
0.69-0.77 in harzburgite and 0.81-0.93 in dunite)
(Fig. 3b) (Ishizuka 1985, 1987). In the Kamuiko-
tan belt, the Takadomari harzburgite body near
Horokanai is similarly highly depleted, but the
Iwanai-dake body contains common harzburgite
and the southernmost Nukabira body consists of
lherzolite (Fume et al. 1997; Tamura et al. 1999)
(Fig. 3b). Tamura et al. (1999) also reported
dunite with highly chromian spinel and magnesian
olivine in the Nukabira lherzolite (Table 2). The
depleted harzburgite of the Horokanai and Taka-
domari complexes has been thrust onto the Ka-
muikotan metamorphic rocks, which include
typical jadeite-bearing blueschists (Ishizuka 1985,
1987; Sakakibara & Ota 1994).
In the Hidaka belt, to the east of the Kamuiko-
tan belt, the Poroshiri ophiolite has a relatively
complete succession (Miyashita 1983). This body
contains abundant troctolite-anorthosite cumu-
lates (Miyashita & Hashimoto 1975), whose oli-
vine-plagioclase crystallization trend closely
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 603
follows that of MORB, and which is thought to be
an L-type ophiolite. Based on geological data,
Miyashita & Yoshida (1988) postulated a Cretac-
eous age for the ophiolites in the Hidaka zone.
The Horoman complex is a well-layered perido-
tite-gabbro body in the southern part of the
Hidaka belt. The main harzburgite-lherzolite
complex has spinel with Cr-number 0.2-0.6 and
contains layers of spinel-rich dunite-wehrlite,
gabbro, and minor banded dunite-harzburgite (spi-
nel Cr-number 0.8-0.9) (Takahashi 1991). The
pervasive occurrence of spinel-pyroxene symplec-
tite after garnet suggests a deep mantle origin.
Rare, large corundum crystals in the metagabbro
suggest that the peridotite body is a fragment of
recycled oceanic crust-mantle, which has been
subducted and then emplaced as a diapir (Morishi-
ta & Arai 2001). L-type residual mantle and
mafic-ultramafic cumulate rocks constitute the
bulk of this peridotite body, but the occurrence of
highly depleted harzburgite indicates that a pre-
existing DH-type ophiolite was also incorporated
into the diapir.
Dobretsov et al. (1994) compared ophiolite-
blueschist belts in central Hokkaido and Sakhalin
and correlated the West and Central (Langeri-
Susunai) Sakhalin Zone with the Sorachi-Yezo
and Kamuikotan belts in Hokkaido. Vysotskiy et
al. (2000) described the boninite-bearing Shelting
ophiolite from central-eastern Sakhalin (Fig. 5).
This ophiolite, composed of dunite-harzburgite,
websterite- orthopyroxenite and gabbronorite
units, occurs as a tectonic slice in fault contact
with Jurassic-Cretaceous volcaniclastic rocks
containing boninite (the Rakitinskaya suite). The
nearby Berezov massif (Fig. 5) has the same
character and lithological sequence. Spinel in the
dunite and harzburgite of this ophiolite is unu-
sually chromian (Cr-number 0.86-0.89) (Fig. 3b,
Table 2). The highly depleted nature of the mantle
peridotite and the presence of the adjacent ortho-
pyroxene-type cumulate rocks help define this
ophiolite as a DH-type. The associated boninite
has 25-30 modal % orthopyroxene phenocrysts,
which show reverse zoning with an iron-rich core
(Mg-number 0.72-0.75) and a magnesian rim
(Mg-number 0.84-0.89). The spinel micropheno-
crysts also show reverse zoning with iron-rich
cores and iron-poor rims (Cr-number 0.80-0.83).
Vysotskiy et al. (2000) suggested that crystal-
lization of the boninitic magma under more and
more reducing conditions was due to the introduc-
tion of a hydrogen-rich fluid.
Ophiolite belts in the Koryak Mountains
The Koryak Mountains extend from northern
Kamchatka to the Bering Straits (Fig. 2). Fujita & Newberry (1982) described the general geology of
this area and Palandzhjan (1986) studied the
ultramafic rocks. Fujita & Newberry (1982) pro-
posed that the Koryak ophiolites were emplaced in
the Jurassic. Pushcharovsky et al. (1988) and
Sokolov (1992) outlined the geological structure
of the Koryak Mountains, in which Early Palaeo-
zoic ophiolites of the Penzhina zone (also called
the Talovsko-Pekulneyskaya zone or the Ust-
Belskaya zone) are thrust over the Koryak nappes,
which contain abundant Mesozoic and rare Pa-
laeozoic ophiolites. Stavsky et al. (1990) proposed
a plate-tectonic model for accretion in the Koryak
Mountains, and Tilman & Bogdanov (1992) pro-
duced a comprehensive geotectonic map of the
area. Dobretsov (1999) compiled chronological
and petrological data from blueschists in this area.
Penzhina zone and Taigonos Peninsula
The Penzhina zone comprises the innermost part
of the Koryak orogenic belt, where ophiolites and
blueschists of various ages are exposed (Fig. 2).
The Pekulney Range in the northern Penzhina
zone has a large Proterozoic(?) metamorphosed
mafic-ultramafic complex and Jurassic ophiolites
(Dobretsov 1999). The Ust-Belaya ophiolite of
Early Palaeozoic age (560Ma K-Ar age on
gabbro) consists of tectonic slices of pillow lava
with chert, dunite-peridotite with gabbro-an-
orthosite layers, and gneissose metagabbro with
eclogitic rocks (Dobretsov 1999). The entire com-
plex was thrust onto the Late Jurassic-Early
Cretaceous accretionary complexes of the Koryak
nappe system (Sokolov 1992). The Ganychalan
('Kharitoninskii' of Dobretsov) nappe in the
southern Penzhina zone also has a complete
ophiolite sequence of Early Palaeozoic age
(430 Ma). The Kuyul ophiolite in the southern
Penzbina zone, mid-Triassic to Jurassic in age, has
a complete succession, whose chemistry and
mineralogy have an SSZ affinity (Khanchuk &
Panchenko 1994; Luchitskaya 1996). The mineral
chemistry of the harzburgite indicates that this is
an H-type ophiolite. According to Dobretsov
(1999), blueschists in the Koryak Mountains
show chronological peaks at 330 Ma (Penzhina),
300 Ma (Penzhina), 180 Ma (Penzhina and Taigo-
nos) and 150 Ma (Penzhina, Pekulnei and m61ange
blocks). Associated ophiolites have ages of 380-
450Ma, 208-290Ma and >150Ma, and the
blueschists of each period are slightly younger
than the adjacent ophiolites.
The Taigonos Peninsula constitutes the south-
ernmost part of the Penzhina Zone (Figs 2 and 6).
Several ophiolite bodies are present along the
eastern coast of the peninsula, such as the Elis-
tratova ophiolite, the Kengeveem metagabbro
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604 A. ISHIWATARI ETAL
P~
m Z
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m
rd
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9
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Z
rO
9 dd~d~d~__ ~ ~ dd~dddd dddd d
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dd~dd~dd
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........ ~ ..... ~.~.
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 605
r.~
0~
0~
r/j
Z
9
9 6~6~ ~ ~ dd~dddd dddd d
Mddd
~d~ ~d~ e~666 d666d~66
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606 A. ISHIWATARI ETAL
N o ~~ S~.~,in J
Ta,ar
' \4"
E142 ~
Quaternary sediments
Tertiary sediments
2~ Cretaceous terrigenous
sedimentary rocks
High-pressure metamorphic rocks
Jurassic-Cretaceous volcaniclastic
rocks, chert, limestone and basalt
Gabbro-norite unit
Websterite-orthopyroxenite unit
~ Dunite-harzburgite unit .X:~ ~' ~\ I Berezov
~7 ~-N50
.~-~.~ Shelting
massif
,1~~\ Seaof
9niya I a.~ N49 ~
E143~ E144~ /
100 km /
\-,-~- ~ Shelting Cape
Fig. 5. Geological map of the DH-type Shelting Cape ophiolite in central eastern Sakhalin (after Vysotskiy et al.
2000). (See Fig. 2 for location.)
complex, the Nablyudeny complex, and the Povor-
otny m61ange. Although the Kengeveem metagab-
bro complex does not have mantle peridotite, the
other three complexes do (Saito et al. 1999;
Bazylev et al. 2001). The Nablyudeny harzburgite
has the most chromian spinel (Cr-number 0.8),
whereas the harzburgites of the Elistratova ophio-
lite have moderately chromian varieties (Cr-
number 0.3-0.7), and the Povorotny lherzolite has
aluminous spinel (Cr-number 0.2) (Table 2). Some
spinels in the Povorotny complex are chromian in
composition with Cr-number up to 0.7 (Bazylev et
al. 2001). Here again, an extreme diversity of
residual mantle peridotite is present among these
ophiolites (Fig. 3c; Table 2).
The Elistratova ophiolite consists of a central
cumulate gabbro body, locally cut by sheeted
dykes, whereas the northern and southern ultra-
mafic bodies are mostly composed of residual
mantle peridotite (Belyi & Akinin 1985; Saito
et al. 1999) (Fig. 6). Disseminated spinel in
peridotite of the southern body is moderately
aluminous (Cr-number 0.30-0.50), whereas that
in the northern body is more chromian (Cr-num-
ber 0.40-0.65) (Fig. 3c). The gabbroic body is intrusive into the southern ultramafic body with a
clear-cut chilled igneous contact. Many dykes,
which contain ultramafic xenoliths, also intrude
into the ultramafic rocks. Saito et al. (1999)
interpreted the northern ultramafic body and the
central gabbroic body as an intact island arc
ophiolite. The southern ultramafic body is less
depleted than the northern one and may have
served as a subvertical wall for the gabbroic
magma chamber. The gabbroic rocks vary upward
from olivine gabbronorite through gabbronorite to
hornblende gabbronorite. The coexistence of ex-
tremely calcic plagioclase (about An90) with rela-
tively iron-rich clinopyroxene (about Mg-number
0.80) clearly indicates an island arc basalt affinity
for the magma (Fig. 4b). Podiform chromitite
from the northern m61ange has a Cr-number of
0.70, and the chromite grains are characterized by
Fe3+-poor cores containing many hydrous inclu-
sions (hornblende and phlogopite), suggesting the
involvement of a hydrous, reducing fluid in podi-
form chromitite formation (Tsujimori et al. 1999).
The Kengeveem metagabbro complex is distinct
from the gabbroic section of the Elistratova
ophiolite. The metagabbros are almost devoid of
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 607
ii: iiiiiiiiiiiiii :iiiii !!!i!iil ii iiiii ! !iiiiiii!iiiiiiiiii! !iiiiiii!ii::ii!!ii!iiiiii!iiiiiiii!iiiiiiiiiiiiiii!!iiii N
9 :.'.5.'.:'." :.:": - "" ":"i:. "J'i:.'.' ",:'~.:',-.: e':3::?:.:.:: I::'.~:'.--W:'.:'-:~ ~ ~.:;.~.:::.'..':-:i:.:;?~:.;::.~.;f ? Elistratova
9 Z.:.?.--"~0~ ''..(-?(ii Peninsula ~ "
~,~..:.:.:~:..:::.::~ / '-,,,, 3o.~,"
/~,~.-.~e ngeveem ~ ~20~'~!t
~...."~ Nablyudeny ~"~35
Cape ..~'-'L~ .... .: ~"~0 Taigonos Povorotny ~ 40
...... .::..: ~, ,~..;:::~:.::.~ --I 50km I _~r~~ N60O
" ~I =====================================================================
E158 ~ E160 ~ E162 ~ .
Elistratova ophiolite Country rocks 3"~L,~~2/ 40
Jurassic chert ~ Tertiary Elistratova
and basalt ~?_.:.i.i~ sediments Peninsula
~1 iabase dyke ~ Lower Cretaceous
complex sediments
Layered gabbro ~ Jurassic
complex andesites
[~ Ultramafic cumulates
~Q Residual mantle peridotites
(North Body and South Body)
Serpentinite melange ~ Layering 0 5 km ,~o/,\~'~
9 i . i i 9 ~o !~o .~o
NW ,~/ff~, ~.~\,~,~'o',' SE
I
Fig. 6. Geological map and schematic cross-section of the Elistratova ophiolite in the Taigonos Peninsula (after Belyi
& Akinin 1985; Saito et al. 1999; and our data). An intact ophiolite sequence composed of the northern ultramafic
body (normal harzburgite) and central gabbro body, with ultramafic cumulates between them, and sheeted diabase
dyke complexes in the gabbro, intrudes into the southern ultramafic body (less depleted harzburgite).
orthopyroxene, which is abundant in the Elistrato-
va gabbros, but contain abundant clinopyroxene
(Mg-number 0.80) and plagioclase (An60) (Fig.
4b), indicating a MORB affinity. On its northem
side, the Kengeveem gabbro body is associated
with Ordovician sedimentary rocks, and thus they
may be significantly older than the Mesozoic
ophiolites of the Taigonos Peninsula. This body
may possibly be the equivalent of the Early
Palaeozoic Ganychalan ophiolite in the Penzhina
zone.
Mainits zone
The Mainits zone in the central Koryak Mountains
is characterized by the occurrence of many ophio-
lites associated with Jurassic-Cretaceous accre-
tionary complexes (Stavsky et al. 1990). Within
an area of 40 km 100 km, there are several
ophiolitic bodies such as the Tamvatney lherzolite
massif, the Yagel serpentinite m61ange, the Kras-
naya Mountain harzburgite nappe, and the Sred- naya Mountain dunite body, all of which are
associated with island arc type gabbro and volca-
nic rocks (Fig. 7). The Tamvatney body contains
mostly lherzolite (spinel Cr-number 0.20-0.30)
(Table 2) with minor harzburgite (spinel Cr-num-
ber 0.35-0.50) (Dmitrenko et al. 1990), and
contains some eclogitic inclusions. The Yagel
serpentinite m61ange consists of blocks of lherzo-
lite, harzburgite (spinel Cr-number 0.4), dunite,
gabbro, picritic sheeted dykes and pillow basalt.
On the other hand, the Krasnaya Mountain ultra-
mafic complex, which occurs as a subhorizontal
nappe thrust onto the Jurassic-Cretaceous accre-
tionary complexes and the Yagel m61ange, contain
highly chromian spinel (Cr-number 0.8) and
highly magnesian olivine (Fo90) (Fig. 3d, Table 2).
The geological map of Dmitrenko et al. (1985)
shows that the southern half of the complex
consists mostly of depleted harzburgite, whereas
the northern half is composed of ultramafic
cumulates and dunite with orthopyroxenite veins
(Fig. 8). The ultramafic cumulate includes iron-
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608 A. ISHIWATARI ETAL
E174 ~ E176 ~
Tamvatney massif ~~ <, ......... :-'J," ' .~"-~
i!!!:: i i i i ii i i
~~ i20kmi!i i!i
.... Post-accretionary terrigenous Jurassic?
sedimentary rocks (Cretaceous) iiilil iii!! i Ophiolitic mafic liiii,iiiii Accreted Jurassic rocks blocks rocks
1: mainly island-arc volcanic rocks, Ophiolitic f il 2. i'i 2: mainlygreywacke ultramafic rocks
Fig. 7. Ophiolitic rock complexes of the Mainits Zone in the Koryak Mountains, NE Russia (simplified from Stavsky
et al. 1990). (See Fig. 2 for location.)
rich harzburgite, dunite, orthopyroxenite and web-
sterite. The Krasnaya complex has a lithological
assemblage typical of the DH-type ophiolite,
although gabbroic and volcanic rocks are missing.
The Serdnaya Mountain dunite body includes
dunite and chromitite. Disseminated spinel in the
dunite is highly chromian (Cr-number 0.85) (Fig.
3d, Table 2). Although the Mainits Zone is a small
area, both L-type and DH-type ophiolites are
present, reflecting wide petrological diversity in
the residual mantle peridotite.
Discussion
Multiple nappe pile of ophiolites of various
ages
Irwin (1981) first described multiple nappe piles
of ophiolites from the Klamath Mountains of
the western USA, where the Jurassic Josephine
ophiolite is tectonically overlain by an Upper
Palaeozoic-Triassic ophiolite, which is in turn
structurally overridden by the Lower Palaeozoic
Trinity ophiolite, with intervening blueschists and
accretionary complexes. Emplacement of older
ophiolites over younger accretionary complexes
and ophiolites is a regular rule not only for the
Klamath Mountains but also for SW Japan (the
Upper Palaeozoic Yakuno ophiolite is tectonically
overlain by the Lower Palaeozoic Oeyama ophio-
lite), NE Japan (the Lower Palaeozoic Miyamori
ophiolite is thrust over the Jurassic accretionary complex and Jurassic-Cretaceous ophiolites in
central Hokkaido), and northeastern Russia (Low-
er Palaeozoic ophiolites of the Penzhina zone are
thrust over the Koryak nappe system containing
abundant Mesozoic ophiolites). The Jurassic(?)
Mikabu ophiolite and the Tertiary Mineoka (-
Setogawa) ophiolite are both present on the Pacific
side of SW Japan, and one of the youngest
ophiolites on Earth is present in East Taiwan
(further review and references have been given by
Ishiwatari (1991, 1994)). This is consistent with
the long-lasting (Phanerozoic) subduction and
accretion of the circum-Pacific orogenic belts.
Petrological diversity of the northwestern
Pacific margin ophiolites
Nicolas & Jackson (1972) demonstrated that lher-
zolite is the dominant mantle peridotite in the
western Mediterranean ophiolites, whereas harz-
burgite dominates in the eastern Mediterranean.
This implies that the mantle composition is rela-
tively homogeneous for more than 1000 km along
the ophiolite belt, although small-scale heteroge-
neities may occur in the boundary area (e.g. in the
Balkan Peninsula). Harzburgite represents residual
mantle left after higher degrees of melting of
lherzolite, and the degree of melting is represented
by various mineralogical parameters such as in-
creasing Cr-number of spinel, increasing Fo con-
tent of olivine, and decreasing A1203 content of
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 609
o I~ --
"= " "- ~='~ 9 ~ x ~ co
_ _ ~- _
,.. ~ ~ 05s o
.-- ,.,,.~, ~ ~ ~.- =~' ~.~~~ =~>~'=~, o o ~=~,-~ o ~ -o ~ ._.-,-.
9 ,~ ~ ~, ~
-- .~ -- r !,4 '-5
13_ ._~ -'o..u ~z:) o~ o 03 [.-...
r./'3
00
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.-o
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o
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's
o
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610 A. ISHIWATARI ETAL
orthopyroxene (Ishiwatari 1985a). DH-type ophio-
lites have not been found in the eastern Mediterra-
nean area.
However, the degree of depletion may vary
significantly within a single 'Tethyan-type' ophio-
lite as in the Bay of Islands (Edwards 1995), and
high-Cr spinel forms chromitite ore in many
ophiolites far from the Pacific Rim, such as in
Albania (Bulqiza massif: (~ina et al. 1987) and the
southern Urals (e.g. Kempirsai massif: Melcher
et al. 1997), where boninitic volcanic rocks with
high-Cr spinel microphenocrysts are also present
(Spadea & Scarrow 2000). In these massifs, the
high-Cr spinel is generally restricted to chromitite
ores, and associated dunites and disseminated
spinel in the surrounding harzburgite are more
aluminous (e.g. Cr-number 0.41 in Kempirsai).
Mantle sections composed totally of highly de-
pleted harzburgite (spinel Cr-number >0.7) are
characteristic of the western Pacific ophiolite
belts. Such harzburgite is associated with ortho-
pyroxene-type cumulates, which may have crystal-
lized from boninitic or island arc tholeiitic melts,
suggesting an SSZ origin of these ophiolites. The
Shelting ophiolite in Sakhalin and the Mt. Kras-
naya ultramafic complex in the Koryak Mountains
are new examples of the DH-type ophiolites.
Previously known DH-type ophiolites include
Horokanai-Yakadomari (Hokkaido; Ishizuka 1985,
1987; Tamura et al. 1999), Papua (England &
Davies 1973; Jaques & Chappell 1980) and
Adamsfield (Tasmania; Varne & Brown 1978).
Highly depleted harzburgite (spinel Cr-number
0.77) and dunite (spinel Cr-number 0.81) were
drilled from a conical serpentinite seamount in the
Mariana forearc. Relatively fertile harzburgite (Cr-
number 0.36) and intermediate varieties were also
recovered from this site (Ishii et al. 1992), and the
spinel Cr-number of all mantle peridotites from
this seamount averages 0.61 (Fig. 9a). Harzburgite
samples from the nearby landward walls of the
Mariana Trench also have Cr-rich spinel (Cr-
number 0.55-0.69) (Bloomer & Hawkins 1983),
and a forearc seamount in the Izu islands has
similar rocks. However, lherzolite samples with
spinel Cr-number of 0.27 were recovered from the
Mariana Trough, an active back-arc rift zone
(Ohara et al. 2002). Mantle peridotites from the
extinct Parece Vela rift zone and southern Mariana
Trench walls near Yap Island contain spinels of
intermediate composition (Cr-number 0.43-0.52)
(Bloomer & Hawkins 1983; Ohara et al. 1996)
(Fig. 9a).
On the other hand, peridotites from fracture
zones of the slow-spreading South Atlantic-
Southwest Indian ridges such as Islas Orcadas,
Vulcan and Bullard have aluminous spinels (Cr-
number 0.15-0.30) (Dick 1989). The Bouvet fracture zone has peridotites with spinel Cr-num-
ber of 0.34-0.55, similar to peridotites of fracture
zones along the fast-spreading East Pacific Rise
(such as the Garrett fracture zone) and the Hess
Deep (Cr-number 0.35-0.45) (Cannat et al. 1990;
Arai et al. 1996; Edwards & Malpas 1996).
Spinels from forearc peridotites are clearly more
Cr rich and more depleted when compared with
those from mid-oceanic ridge peridotites. Boni-
nites from the Mariana and Tonga forearcs have
highly chromian spinel (Cr-number 0.70-0.90)
and are believed to have formed by hydrous
melting of a depleted mantle (Bloomer & Haw-
kins 1987; Sobolev & Danyushevsky 1994). Pri-
mitive magnesian andesite, containing chromian
spinel (Cr-number >0.74) and magnesian ortho-
pyroxene (Mg-number 0.88), has also been re-
ported from Oligo-Miocene island arc volcanic
rocks of Japan in association with tholeiitic basalt
and calc-alkaline andesite (L6pez & Ishiwatari
2002). The occurrence of diverse mantle perido-
tites including highly depleted harzburgite (spinel
Cr-number >0.70) in the Mariana forearc suggests
that the ophiolite belts of the northwestern Pacific
margins also originated in intra-oceanic SSZ
environments extending from landward trench
walls to back-arc basins.
It should be noted, however, that the variation
in Os isotopic composition among ophiolitic chro-
mites, including those from the northwestern
Pacific margins, is significantly less than has been
reported for oceanic peridotites and MORB. Thus,
there is little evidence to suggest modification of
the mantle's original Os isotopic composition via
radiogenic melts or fluids derived from subducting
slabs (Walker et al. 2002). This may partly reflect
the extremely low concentration of platinum-
group elements in such fluids but essentially sug-
gests that simple, high-temperature melting of
homogeneous, depleted MORB mantle (DMM) is
the major ophiolite-forming process.
Occurrence of the ophiolites with thick
oceanic crust
Granulite-facies, two-pyroxene spinel metagabbro
(sometimes with garnet) occurs at the Moho level
of some circum-Pacific ophiolites in Japan, eastern
Russia and Alaska (Ishiwatari 1985b; DeBari &
Coleman 1989; Vysotskiy 1994; Tsujimori &
Ishiwatari 2002). This observation suggests a
relatively thick oceanic crust for these ophiolites,
which are believed to have formed in island arc
(DeBari & Coleman 1989), marginal basin (Ishi-
watari 1985b), and/or oceanic plateau (Isozaki
1997) environments. Granulite-facies metagabbros
(mostly spinel and garnet free) have also been
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 611
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612 A. ISHIWATARI ETAL
recovered from mid-oceanic ridges, but Gaggero
& Cortesogno (1997) estimated that the pressure
of formation for these rocks was 0.3 GPa (c.
10kin depth). This is much lower than the
pressure (0.5-1.1 GPa) estimated for ophiolitic
granulite-facies metagabbros (Ishiwatari 1985b;
DeBari & Coleman 1989; Vysotskiy 1994; Tsuji-
mori & Ishiwatari 2002).
Coffin & Eldholm (2001) pointed out that
sections of oceanic crust in large igneous pro-
vinces (LIPs) are two or five times thicker than
those of 'normal' oceanic crust and postulated that
some ophiolites are of LIP origin. Xenoliths of
granulite-facies, two-pyroxene spinel metagabbro
have been reported in alkali basalt from the
Kerguelen Archipelago, a typical LIP (Gr~goire
et al. 1998). However, analogous spinel metagab-
bro xenoliths are also known from continental
margins (e.g. Francis 1976), although higher-pres-
sure garnet granulites may comprise typical lower
continental crust (e.g. Loock et al. 1990). Based
on their petrology, mineralogy and geochemistry,
circum-Pacific ophiolites with relatively thick ma-
fic crust represent SSZ lithosphere (DeBari &
Coleman 1989; Ishiwatari et al. 1990). Many are
tectonically underlain by blueschist and younger
accretionary complexes, suggesting that they re-
present the hanging wall of the subduction zone
(i.e. mantle wedge and overlying crust). The Os
isotope character of chromitite in the Yakuno
ophiolite does not support an origin in a super-
plume-related LIE which would be characterized
by an isotopically distinct HIMU or enriched
mantle (Walker et al. 2002).
Origin of the ophiolite-blueschist
assemblage
Japanese ophiolites commonly have metamorphic
soles composed of blueschist and are tectonically
underlain by younger, sediment-rich accretionary
complexes, which contain greenstones of OIB and
MORB origin (Table 1). For example, in south-
western Japan, the Ordovician Oeyarna ophiolite
(>450 Ma) is underlain by the 320 Ma Renge
blueschist and the Upper Permian (250Ma)
Akiyoshi accretionary complex. This spatial rela-
tionship suggests tectonic erosion or non-accretion
during the intervening Siluro-Devonian time. A
similar gap exists between the Lower Permian
(280 Ma) Yakuno ophiolite and the underlying
Jurassic Tamba accretionary complex (150Ma).
The accretionary complex is characterized by
'oceanic plate stratigraphy' composed of green-
stone, chert, limestone, mudstone and sandstone in
a younging order (Isozaki 1997). The basal green-
stone commonly includes tholeiitic and alkaline seamount basalt (OIB) with high Ti and Nb
concentrations, but the ophiolite itself is almost
free of OIB.
In the present-day western Pacific, the Izu-
Mariana and Tonga subduction zone environments
are characterized by the presence of ophiolite
outcrops on the trench slopes (Bloomer & Haw-
kins 1983; Bloomer & Fisher 1987) and blues-
chists (Maekawa et al. 1993, 1995), by the
absence or scarcity of accretionary complexes,
and by the currently active back-arc spreading. In
contrast, areas off northeastern Honshu and Hok-
kaido are characterized by the development of vast
accretionary complexes (Taira 1985) without sub-
marine ophiolite or blueschist outcrops and with-
out active back-arc spreading (Fig. 9). These
different environments may have been repeated in
any segment of the Japanese orogenic belt
throughout Phanerozoic time. Periods of oceanic
island arc and marginal basin development (ophio-
lite formation) and tectonic erosion (blueschist
metamorphism) might have alternated with peri-
ods of normal subduction, during which accretion-
ary complexes were developed.
The ophiolite-blueschist association is well
documented in the Japan-Primorye area, e.g. the
Oeyama ophiolite-Renge blueschist (Tsujimori &
Itaya 1999) and Sergeevka ophiolite-Shaiginskiy
blueschist (Kovalenko & Khanchuk 1991; Zakhar-
ov et al. 1992). In the NE Japan-Sakhalin belt,
the Palaeozoic Miyamori ophiolite-Motai blues-
chist pair (Maekawa 1988; Ozawa 1988, 1994)
and the Mesozoic Horokanai ophiolite-Kamuiko-
tan blueschist pair (Ishizuka 1985, 1987; Sakaki-
bara & Ota 1994) are well documented. Although
a major blueschist belt is absent in the Koryak
Mountains, many blueschist blocks occur in the
Palaeozoic and Mesozoic accretionary complexes
(Stavsky et al. 1990; Dobretsov 1999).
It is likely that periods of accretion and non-
accretion, as represented by the present-day Nan-
kai Trough and Mariana Trench, respectively, have
been repeated many times in different segments of
the Japan-NE Russia accretionary orogenic belts
in the past. Periods of ophiolite-blueschist forma-
tion and tectonic erosion at subduction zones
might have been followed by periods of massive
accretion. Tectonic underplating of accreted sedi-
ments beneath the mantle wedge might have
facilitated the uplift of overlying ophiolite-blues-
chist assembIages. This idea is compatible with
the geochemical signatures of ophiolitic rocks
showing SSZ affinities.
Conclusions
The northwestern Pacific margin extending from
Japan to Russia has many ophiolites of widely
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OPHIOLITES IN JAPAN AND FAR EAST RUSSIA 613
varying ages, different petrological characteristics
and distinctive tectonic histories. The following
geological features suggest that these ophiolites
probably formed in island arc environments in
intra-oceanic settings: extremely diverse degree of
melting in the residual mantle peridotite up to
clinopyroxene disappearance and spinel Cr-num-
ber >0.70; the common occurrence of hydrous
minerals and various metasomatic features in the
mantle section; the common association with
blueschist rocks; the presence of unusually thick
oceanic crust. The modem Mariana and Tonga
trenches, where ophiolitic rocks including highly
depleted harzburgite and typical blueschist have
been dredged from the sea floor, may be the
modem analogues. The orogenic belts from Japan
to NE Russia may have evolved through repeated
stages of non-accretion, in which SSZ ophiolites
and blueschists formed, and accretion, in which
accretionary complexes mainly composed of clas-
tic and volcaniclastic rocks developed.
The association of highly depleted mantle harz-
burgite and orthopyroxene-type cumulate rocks is
reinforced by reported occurrences of DH-type
ophiolites from NE Russia (Shelting and Kras-
naya). These ophiolites have only been reported so
far from the western Pacific margins such as
Hokkaido (Horokanai), Papua, and Tasmania
(Adamsfield). The association of some DH-type
ophiolites with boninitic volcanic rocks (Shelting,
Papua, and possibly Mariana and Tonga) suggests
that the depleted harzburgite is a residuum after
boninitic melt production, although boninite is
also reported from some ophiolites with less
depleted peridotite (e.g. Robinson et al. 1983;
Spadea & Scarrow 2000). Some primitive island
arc tholeiite and magnesian andesite magmas
could also coexist with the depleted harzburgite.
The depleted harzburgite may form by either high-
temperature dry melting of primary mantle or
hydrous melting of previously depleted mantle.
However, Os isotope studies of ophiolitic chromi-
tites do not support much involvement of slab-
derived fluids in mantle melting. The Os isotope
data are also inconsistent with an oceanic plateau
(or LIP) origin postulated for some ophiolites with
thick crustal sections. These instead may represent
robust magmatic activity in SSZ environments,
where they would be associated with highly
depleted harzburgite massifs.
We are grateful for T. Tsujimori, D. Saito, S. Miyashita,
A. P. Stavsky, O. Morozov, S.A. Shcheka, A. I. Khan-
chuk, S. G. Byalobzhesky, W. B. Bryan and J. Hourigan
for their help with fieldwork in Far East Russia. A.I.
thanks S. A. Palandzhjan for valuable information on the
ophiolites of the Koryak Mountains, and Y. Dilek for his
encouragement in writing this paper. S. Maruyama is
acknowledged for his help in arranging fnancial support for our fieldwork. A.I. also acknowledges Grant-in-Aid
for Scientific Research (C)-(2)-10640462 and -14540447
by the Ministry of Education, Japan. S. Arai is thanked
for several discussions. Y. Furuhashi and D. Saito
contributed the mineral analyses.
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Water Air Soil Pollut (2016) 227: 385 ssMarl DOI 10.1007/s11270-016-3089-3 Geochemical Processes Controlling Fluoride Enrichment in Groundwater at the Western Part of Kumamoto Area, Japan Shahadat Hossain · Takahiro Hosono · Heejun Yang · Jun Shimada Received: 14 December 2015 /Accepted: 13 September 2016 /Published online: 27 September 2016 Springer International Publishing Switzerland 2016 Abstract This paper first time reports the geochemical a significant role in fluoride mobilization. Weakly alka- processes that are controlling fluoride enrichment in the line nature of groundwater with high pH (7.05-9.45) groundwater of western Kumamoto area, Japan. Fifty expedites the leaching process of exchangeable F from (50) groundwater samples were collected and analyzed F-bearing minerals as well as favors desorption of F for the study where fluoride (F） concentration ranges from metal oxide surfaces. High HCO3 and high PO4 from 0.1 to 1.57 mg/L. About 58 % of the shallow in the groundwater facilitate desorption process as groundwater and 26 % of the deep groundwater sam- competing anions, while high Na+/Ca2+ ratio largely ples contain fluoride concentration beyond the Japanese control this process by decreasing positive-charge den- drinking water permissible limit (0.8 mg/L). High F is sity of the metal-oxide surfaces. High Na+/Ca?+ ratio is largely accumulated in the stagnant zone of the attributed due to the cation-exchange process, while Kumamoto Plain area and associated with Na-HCO3- high pH and HCO3 are the result of both silicate type groundwater. High pH, high HCO3, low Ca2+, and hydrolysis and microbial reduction processes. In addi high Na+ are the major characteristics of high-F tion, calcite and fluorite seem to have a control on groundwater. Hydrolysis of F -bearing minerals and groundwater fluoride geochemistry. desorption of F from hydrous metal oxides are consid- ered to be the primary sources of fluoride in ground- Keywords Fluoride · Geochemistry · Alkaline water. A positive correlation between F and Na+/Ca2+ groundwater· Cationexchange · Kumamoto basin · Japan ratio (r2 = 0.53) indicates that major ion chemistry plays Electronic supplementary material The online version of this 1 Introduction article (doi:10.1007/s11270-016-3089-3) contains supplementary material, which is available to authorized users. Presence of elevated level of toxic elements in S. Hossain () · H. Yang · J. Shimada groundwater is a serious issue for water supplies Graduate School of Science and Technology (GSST), Kumamoto around the world as it impacts on human health. University, 2-39-1 Kurokami, Kumamoto-Shi 860-8555, Japan Considering toxicity, high fluoride (F) in ground- e-mail: sahedmc @ yahoo.com water has been receiving worldwide attention long S. Hossain ago since Dean and Elvove (1937) established the Atomic Energy Center, Chittagong, Bangladesh Atomic Energy relationship between fluorosis and fluoride in drink- Commission, Chittagong 4209, Bangladesh ing water. Of few other sources, drinking water is often the major source of human intake of F-, where T. Hosono Priority Organization for Innovation and Excellence, Kumamoto people essentially rely on groundwater for their University, 2-39-1 Kurokami, Kumamoto-Shi 860-8555, Japan drinking purpose (Cao et al. 2000; Edmunds and Springer 385Page 2 of 14 Water Air Soil Pollut (2016) 227: 385 Smedley 2013).Although F- concentrations in phosphates (PO43) in marine sediments (Edmunds and drinking water of around 1 mg/L are considered Smedley 2013). favorable by many countries, different countries fol- Numerous researchers have investigated the fluoride low different guideline values. For example, most contamination and subsequent rock-water interactions European countries, Canada, India, etc., use WHO in various aquifers with diverse geology worldwide (Li (World Health Organization 2011） value,i.e., et al. 2015, and references therein). It is concluded that 1.5 mg/L as national standard, whereas China set not only the sources but also the geochemical processes the value as 1.0 mg/L, and in Japan, maximum F as well as climate play a major role in controlling consumption limit in drinking water is set to 0.8 mg/ fluoride concentrations in groundwater system. For ex- L (Ministry of Health, Labor and Welfare, Japan; ample, high pH and soil salinity are highlighted as the http://www.mhlw.go.jp/english/policy/health/water_ key factors that controlled high-F groundwaters in an supply/4.html). Lower concentration of F- is arid sedimentary aquifer of China (Fuhong and Shuquin beneficial to the dental health, while higher 1988). Similarly, both Handa (1975) and Dissanayake concentrations are considered detrimental (WHO (1991) discussed the climatic and geologic controls on 2004; 2011). Endemic fluorosis due to intake of high-fluoride groundwater and implied that high fluo- F -rich groundwater is reported in more than 25 ride occurred in arid to semi-arid zones. In turn, countries such as India, China, Korea, Mexico, Sri Edmunds et al. (1989) reported that high-fluoride Lanka, USA, and Eastern African countries (Gaciri groundwater in a carbonate aquifer is controlled by a and Davies 1993; Diaz-Barriga et al. 1997; Wuyi combination of fluorine source (CaF2), longer residence time, and Nat-Ca2+ ion exchange process. A number of Majumdar 2011,and references therein). It is esti- sedimentary basins are found with high-fluoride con- mated that more than 2o0 million people around the centrations along the groundwater flow line with in- world are suffering of chronic endemic fluorosis as a creasing residence time (Travi 1993). Generally, alka- line pH, low Ca2t, high Na, and Na-HCO3s-type consequence of taking high-F- groundwater (Ayoob and Gupta 2006). groundwater are the common characteristics of high- Fluorine is the 13th most common element in Earth's fluoride groundwater in many aquifers (Guo et al. crust (625 mg/kg), and groundwater fluoride primarily 2007; Chae et al. 2007; Rafique et al. 2009; Currell occurred through weathering from their geological strata, et al. 2011; Su et al. 2013). where F -bearing minerals are abundant as well as geo- In this study, for the first time, we report the thermal sources (Gizaw 1996; Carrillo-Rivera et al. potential hydrogeochemical processes that are respon- 2002). Thus, fluorine occurrence is source dependent, sible for the high-fluoride groundwater in Kumamoto mostly found in acidic igneous rocks, mineralized veins, area. Kumamoto is the largest urban groundwater user and sedimentary formations, where biogeochemical re- city in Japan, where about 1 million people complete- actions may take place (Edmunds and Smedley 2013). In ly depend on groundwater for their all-purpose use. It granitic terrains, dissolution of micas (biotite) is the only city in Japan which uses 100 % ground- (K2(Mg,Fe)4(Fe,A1)2[Si6Al2O2o](OH)2(F,C1)2 and am- water in each tap water supply. Thus, groundwater pｈibｏｌeｓ(ｈｏｒｎbｌeｎde）(K,Nａ)o- quality is a major concern to the city inhabitants as 1(Ca,Na,Fe,Mg)2(Mg,Fe,A1)s(Al,Si)8O22(OH)2, where well as to the water supply managers and policy F substitutes for OH positions in their mineral structure, makers. Groundwater with high-fluoride concentra- is the major source of groundwater fluoride (Nordstrom tions up to 4.5 mg/L had been detected locally in et al. 1989; Datta et al. 1996; Chae et al. 2006). However, the western part of Kumamoto basin (Tsuru et al. fluorite (CaF2), the main fluorine mineral, also appeared 2006). However, detailed scientific investigation has as a dominant source of fluoride in some granitic regions never been carried out hitherto to understand the (Handa 1975; Deshmukh et al. 1995). Another important fluoride occurrence in this area. Therefore; present sources of fluorine are apatite (Cas(PO4)3(F,Cl,OH)) and study aims at understanding the geochemical process- fluorapatite (Cas(PO4)3F), which may form in both high- es that control the fluoride enrichment in this area, and low-temperature volcanic areas. Apart from this, which could be helpful to the policy makers in fluorine may accumulate both by adsorption onto clays Kumamoto area as well as scientific arenas of other and also by biogeochemical processes that replace regions with similar hydrogeological conditions. Springer Water Air Soil Pollut (2016) 227: 385 Page 3 of14385 2 Materials and Methods Groundwater area is lying between Shira River water- shed to the north and Midory River watershed to the 2.1 Study Area Settings south. Topographical inclination occurs with gentle slope from north-east to the south-west direction. Kumamoto groundwater area is a Quaternary volcanic P-T path with isothermal decompression at the early phase of area in central Kyushu Island of southern Japan and tional isopleths of Si in phengite and Xmg of chlorite are shown. Correlation of the modelled isopleths allowed the P-T conditions of the prograde and the peak covers an area of nearly 1000 km² (Fig. 1). The study "Kumamoto Plain,” a widely spread plain area (<8 m) area is formed among mountains, highlands, and low- that faces to the Ariake Sea at the south-western part of lands. Aso Mountain (1592 m), the largest active caldera + paragonite + quartz + hematite, epidote + phengite + glaucophane + chlorite + lawsonite + paragonite + quartz + hematite and phengite + glaucophane + Fig. 9 P-T pseudosection of the Wakasa blueschist (WBS 1-2), showing phase assemblage fields with the bulk composition in mol.%. Calculated composi- climate with a significant temperature variations between Kimpo Mountain (665 m) set the west margin. Three summer and winter. Average annual temperature and major highlands with altitude of 100-200 m surrounded precipitation are 16.9 °C and 1986 mm, respectively, the study area at north-eastern and south-eastern sides, during the last 35 years (Japan Meteorological Agency namely, Ueki, Kikuchi, and Takayubaru highlands. data; http://www.jma.go.jp/jma/menu/report.html). Most Kumamoto D 300km KyushuIsland Ueki upland Mt.ASO upland R.Shira 30m Takayubart AriakeSea 20m Ezu Lake 1st aquifer F (mg/L) Elevation (m) 2nd aquifer 0S-0 ·0.0-0.2 Publisheddata 50-100 100-300 0.2-0.4 300-500 Groundwaterwatershed 0.4-0.8 500-800 ofdeepaquifer 800-1200 km 0.8-1.5 1200-2000 0 10 2000-3200 >1.5 3200< Fig. 1 Location of the study area showing major geographic the groundwater potential (m, asl). Hydrogeologic sections along features, groundwater region, and sampling points along major chlorite + lawsonite + paragonite + quartz + hematite stability fields. The modal abundance of paragonite and lawsonite are also shown. Bulkcompositionin mol.%,NCKFMASHO model system(+H,O,Qz) from Tsuru et al. (2006) as for reference along the arrows A-A' and B-B'. The solid contour lines represent Springer 385Page 4 of 14 Water Air Soil Pollut (2016) 227: 385 rainfall occurs between June and July during summer through “paddy field ponding,” where lacustrine sedi- monsoon,which accounts for ~40 % of the total ment clay layer is absent in between the aquifers. precipitation received in Kumamoto area. Groundwater mainly discharges via springs, groundwa- Summaries of the geology, hydrogeology, and ter abstraction, and evaporation process. There are two groundwater evolution processes with substantial hy- major flow lines in the study area: (i) A-A' flow line, where groundwater recharge occurs at Kikuchi highland (2013, 2016). The basement outcrops are Paleozoic to and it surroundings including the eastern foot ofMt. Aso Mesozoic metamorphic rocks mostly distributed in the and mid-stream of Shira River area and laterally flow southeastern and northwest-western parts of the study towards south-west direction and discharge at Ezu Lake area. Cretaceous sedimentary rocks, Tertiary- area. After Ezu Lake, A-A' flow line becomes weak and Quaternary pre-Aso volcanic rocks, Quaternary Aso reach to the Kumamoto Plain area with very low poten- volcanic rocks, and Holocene alluvium deposits are tiality (Supplementary figure), and (i) the other one is superimposed on to the basement (surface geology the B-B' flow line, in which groundwater recharge at the map in Supplementary figure) (Watanabe 1978; Hunter Ueki highlands area and flows towards south-westward 1998). Pre-Aso volcanic rocks are mainly intermediate direction via Kumamoto Plain area up to the coastal to acidic lavas and tuff breccia, which present around zone (Fig. 2). These two flow lines are not connected Mt. Kimpo. Additionally, Aso volcanic rocks are depos- hydrogeologically, but flow characters are almost simi- ited over andesite and basement rocks at the eastern side lar and controlled by the regional geology and topogra- of Mt. Kimpo (Ono and Watanabe 1985). These Aso phy (Hosono et al. 2013). Since the pyroclastic flows are volcanic rocks are characterized into four units as Aso-1 highly porous and permeable, groundwater flow is com- to 4 according to eruption cycle that occurred from older paratively faster along the flow lines especially in A-A' to younger age and collectively known as the Aso flow line with the mean velocity of 40 m day-', and pyroclastic flow deposits (APFDs). Two impervious hydraulic conductivities both in the first and second layers composed of alluvial and lacustrine sands and aquifers are 1.0 x 10-7~1.0× 10-2 and 1.0 × 10-7~1.0 × 10-4 ms-1, respectively (Kumamoto silts are deposited in between Aso-1 and 2 and Aso-3 and 4 units (Kumamoto Prefecture and Kumamoto City Prefecture and Kumamoto City 1995). On the other 1995). In addition to this, a thick marine clay layer hand, Kumamoto Plain area is a stagnant zone due to composed of alternative silt and sand bed of its sluggish nature and groundwater potential is very low Shimabara gravel formation of Pleistocene is placed compare to the two major flow lines (Fig. 1). over Aso-4 unit at Kumamoto Plain area (Fig. 2). Groundwater residence time along the flow lines is The Quaternary Aso pyroclastic flows and alluvium nearly 30 years, whereas in Kumamoto Plain area, it is sedimentary deposits consist the most important more than 55 years (Kagabu et al. 2013). groundwater aquifer system in this area, which is verti- Petrographic study of this area has been carried out o ai o r, q s by several researchers and found that most common aquifers. The upper or first aquifer is basically uncon- rock type is pyroxene andesite. Pre-Aso volcanic rocks fined with shallow depth generally from 0 to 50 m and Aso volcanic rocks both have a wide chemical range below the surface and comprised of partly Quaternary between basaltic to rhyolitic compositions (Ono and Aso-4 pyroclastic and Holocene alluvium deposits. On Watanabe 1985; Miyoshi et al. 2009). Silicate and the other hand, second aquifer or deep aquifer is com- alumino-silicate minerals are the most reported rock- posed of entirely Aso-1 to 3 pyroclastic units with the forming minerals in the volcanic rocks of this area. buried depth of 50-200 m. However, in some regions, Mineralogical composition of APFD shows that plagio- aquifer depths vary depending on the geology and to- clase, clinopyroxene, orthopyroxene, hornblende, bio- pography especially in Kumamoto Plain area, where tite, titanomagnetite (magnetite and ilmenite), and small first aquifer reaches more than 60 m deep (Fig. 2). Groundwater is recharged vertically through meteoric phases; however, the most variable component is infiltration at the highland areas and laterally along the groundmass, which varies significantly in Aso-1 to mountain front. Apart from this, a significant amount of Aso-4 from 45 to 80 % (Watanabe 1978; Hunter groundwater is also recharged to the deep aquifer direct- 1998). On the other hand, kaolinite and smectite are ly in the mid-stream area of Shira River artificially the most stable secondary minerals along with some Springer Water Air Soil Pollut (2016) 227: 385 Page 5 of14385 B Basement B' RechargeZone uplifting Lateral flowtostagnantzone zone (m) F (m) E 200 200 Ueki Lacustrinesediments(Aquiclude) +++.. 100 100 As0-4 Shira R Alluvium Marine clays 0 As0-1~3 0 -100 Hills-plain border zone 100 Aso-2flow lavas -200 Hydrologic basement 200 (Paleozoic basement and Pre-Aso volcanic rocks ) -300 -300 c C' Kumamoto plain 10 10 0 0 sandy silt Sandy loam -20 -20 Ariake Clay -40 Shimabara deposit -40 Sandy gravel -60 -60 Aso-4 -80 -80 Aso 4/3 sediment Aso 2/1 sediment -100 -100 -120 -120 Aso-3 -140 -140 Fig. 2 Simplified hydrogeological cross sections along B-B' and C-C' lines in Fig. 1 other alteration products (Kumamoto Prefecture managed by ministry of land, infrastructure and trans- Geological Map Compilation Committee Report 2008). port of Japan government, Kumamoto prefectural gov- ernment, and Kumamoto city government. Few samples 2.2 Sampling and Laboratory Analysis were collected from private wells, which are mostly used for domestic/agricultural purposes owned by local A total of 50 groundwater samples were collected from people. Three groundwater samples were also collected Kumamoto groundwater area covering both shallow from A-A' flow line to compare with the existed data (10-60 m) and deep aquifer (50-155 m) during and reference for the present study. Groundwater tem- perature, pH, EC, dissolved oxygen (DO), and October-November 2013 (Fig. 1). The sampling was planned considering previous results by Tsuru et al. oxidation-reduction potentials (later converted to pe) (2006), in which groundwater quality investigations were measured in the field site with potable meter were carried out mainly in the Shira River watershed (HORIBA D-54, Japan) using calibrated electrode for area along A-A' flow line and parts of Kumamoto Plain each parameter under minimum atmospheric contact as area. By observing their results, we were convinced to described in Hossain et al. (2016). Acid-washed (10 % look into different flow line, B-B', and entire Kumamoto HNO3, Wako Chemical Ind. Ltd., Japan) high-density Plain area and thus collected 47 groundwater samples polyethylene bottles were used for sample storage. mostly from observation and monitoring wells, Samples were collected after purging either 30 min or Springer 385 Page 6 of 14 Water Air Soil Pollut (2016) 227: 385 three times of well volume to ensure the fresh flowing deep aquifer groundwater change almost similarly along water and filtered with 0.2-μm cellulose-acetate filters each flow line. Figure 3 shows the hydrochemical char- (Advantec?, Japan). Cation and trace metal samples acteristics and evolution of groundwater collected from were stored in a 100-mL bottle with addition of 6 N B-B' flow line and Kumamoto Plain area. Groundwater HNO3 immediately after sampling to make pH< 2, in facies can be categorized into the following four types: order to prevent the depositions of metals. Similarly, Ca-HCO3, Ca-Na-HCO3, Na-HCO3 (mixed cation), and samples for the analysis of alkalinity, anion, and dis- Na-Cl type. Recharge waters are basically Ca-HCO3 solved silica (SiO2) were collected into another separate type in B-B' flow line, which became Na-Ca-HCO3 type 100-mL bottle without preservation. Alkalinity titration within short distance of recharge area and followed by was performed within the day of sampling by volumet- Na-HCO3 type in the middle part of the flow zone and ric titration using 0.05 N H2SO4 and bromocresol green finally appeared as Na-Cl type at the end of B-B' flow methyl red indicator. Major cations (Na*, K+, Ca²+, and near the coastal area. However, this Na-Cl-type ground- Mg²+) and anions (CI, F, NO3, SO4²-, and PO43-) water contains high pH, CI , SO4-, and HCO3 values, were determined by ion chromatography (Compact IC, which revealed that coastal zone waters mostly influ- 761, Metrohm, Switzerland) using SHODEX HPLC enced by marine sediments rather than groundwater column SI-90 4E and SI-90G (Showa Denko K.K flow line. Kumamoto Plain area is greatly dominated Japan） at Hydrology Laboratory, Kumamoto by Na+ and HCOs ions and became Na-HCOs type University. Trace elements such as total arsenic (As) (see Supplementary figure). In addition, A-A' flow line were measured by an inductively coupled plasma-mass groundwater is dominant by Ca2+ and Mg2+ with low spectrometer (ICP-MS) (NexION300D, PerkinElmer, TDS value (<150 mg/L). Ca-HCO is also the major USA) using standard mode and kinetic energy disper- water type in the recharge area of this flow line. It sion (KED) mode at the same laboratory with a detec- changes to Ca-Mg-Na-HCO3 type in lateral flow zone tion limit of 1 μg/L. Precisions were better than ±5 % for of A-A' and finally evolves as Na-HCOs type at the end both major ions and trace elements, evaluated from of A-A' flow line (Hossain et al. 2013). Groundwater pH repeated measurement of standard reference solutions values vary from 7.05 to 9.45 with average values 7.83 with certified concentrations. A microplate spectropho- and 7.77 in first and second aquifers, respectively tometer (MultiskanTM GO, Thermo Fisher Scientific, (Table 1). Total dissolved solids (TDSs) ranged from USA) was used to measure SiO2 concentration by mo- 94 to 3608 mg/L. First-aquifer samples are generally lybdenum ammonium colorimetric method at the wave- more saline than from second-aquifer samples. Higher length of 410 nm. Ion-balance errors (IBEs) were used NO3 concentrations were found in some recharge wa- to evaluate the analytical accuracy, and it was better than ters due to anthropogenic contamination mainly from ±5 % for all the samples. PHREEQC computer code farmland and agricultural activities. On the other hand, with WATEQ4F database (version 2.14.3; Parkhurst and Appelo 1999) was used to calculate mineral saturation indices (SIs) and aqueous speciation of the elements. 2nd aquifer ）1staquifer 3 Results C 3.1 Groundwater Geochemistry Details of the hydrogeochemical evolution of ground- water along the two major flow lines in Kumamoto area are presented in Hossain et al. (2013), and gist informa- tion of the study are mentioned here. The study revealed O that modern precipitation is the main source of 12- Kumamoto groundwater, which rapidly infiltrates with- 1.5 -1.2 -0.9 0.6 -0.3 0.0 0.3 0.6 out significant evaporation through volcanic porous li- CAI 2 (meq) thology. Chemical characteristics of both shallow and Fig. 3 Scatter plot of CAI 1 vs CAI 2 of groundwater samples Springer Water Air Soil Pollut (2016) 227: 385 Page 7 of14385 0.54 15.2 9 Median 3. 3 9 8 Average 57. 1 3 Maximum 15%6. 6 .74 Minimum 4018410 8400480492042 8. 4. Second aquifer 0.62 Number Table 1 Statistical summary of chemical compositions of groundwater samples collected from Kumamoto area 334 Standard deviation G 003 9. 4 3 3 Median 忆州019 7. 2. 1.5 4 2.27 02 Average 6切 6 38 10 Maximum 3 8 5 12 . 00 Minimum 28. 4 0.66 10 First aquifer Number 19 9 19 19 19 191919919199191919191919919 1919 Units mg/L mg/L mg/L mg/L mg/L meq/L meq/L mg/L T8u fluorapatite Parameters +O/+N fluorite Depth E p比m S S Springer 385Page 8 of 14 Water Air Soil Pollut (2016) 227: 385 groundwater in Kumamoto Plain area contains high that contain clay minerals. The secondary minerals such SO42- concentrations, which are mainly of marine ori- as kaolinite and smectite that produced from the gin (Hosono et al. 2013). In Kumamoto basin, high weathering reactions through precipitation as well as phosphate (PO4') and arsenic were also detected in marine and non-marine sediments provide abundant groundwater samples from Kumamoto Plain area, exchange sites to facilitate this process. Moreover, lon- which ranged between 0.48 and 12 mg/L and 0.1 and ger residence time or stagnant nature of groundwater in 60.6 μg/L, respectively. Being a volcanic terrain, hydro- Kumamoto Plain area is favorable for the cation ex- lysis of silicate minerals is found the major weathering change process (Hossain et al. 2013). The major differ- processes in the study area. Changes of groundwater ence in both flow lines is in their redox condition. A-A' chemistry along the flow lines are mostly controlled flow line was dominant with oxic environment without by the weathering of plagioclase minerals with a signif- observing any major redox process. However, anaerobic icant contribution from biotite and pyroxene minerals. biogeochemical reduction processes were predominant Kaolinite and smectite are found most stable secondary in B-B' flow line with parallel increasing of pH along minerals (Hossain et al. 2013). Although most ground- the flow line. Therefore, B-B' flow line is highly redox waters at recharge area had high concentrations of Ca2+ active and concentrations of all redox sensitive elements and Mg2+, Na concentration increased relatively along are mainly controlled by the successive redox reactions. Kumamoto Plain area was classified as heterogenic Nat/Ca2+ ratio from 0.4 to 27 (see supplementary redox condition predominating with Fe(Il)-SO4 reduc- Table A1). For most of the cases, such change in Nat/ ing zone (Hossain et al. 2016). Ca2+ ratio is not accompanied with total cation/Cl equivalent ratio, which indicates that increase of Na+ 3.2 Concentrations and Spatial Variations along the down gradient is the result of other process of Groundwater F apart from mineral weathering only. The study further revealed that cation exchange of Ca?+/Mg?+ to Na+/K+ F concentrations in the groundwater range from 0.12 to is the dominant process along the B-B' flow line from 1.57 mg/L in the first aquifer and 0.10 to 1.17 mg/L in recharge to the discharge area including Kumamoto the second aquifer with mean values of 0.83 and Plain area. Cation exchange processes could be support- 0.55 mg/L, respectively (Table 1). PHREEQC specia- ed by the chloro-alkaline indices CAI 1 and CAI 2 tion calculation shows that fluoride in groundwater is (Schoeller 1967). These chloro-alkaline indices were present as F, MgF+, CaF+, and NaF including some calculated using following equations: other minor species, where F species accounts for 96-- 100 % of the total F. About 58 % wells from the first CAI 1 = [CI-(Na+ + K+)] C1 aquifer and 26 % wells from the second aquifer cross the CAI2 = [CI- (Na+ + K+)]/(SO4²- + HCO3-+ CO²- + NO3) Japanese groundwater standard limit for F (0.8 mg/L). Maximum F concentration, 1.57 mg/L, which crosses The positive values of both CAI 1 and CAI 2 mean WHO limit also (1.50 mg/L), was detected at a depth of that dissolved Na+ and K+ in the water solution are 54 m. The spatial distributions of F show that fluoride exchanging with the Mg2+ and Ca2+ of the clay or rock concentration is generally increasing along the ground- matrix. In turn, if the exchange occurred in the reverse water flow line (Fig. 1). Nonetheless, two samples from order, then the indices would be negative. The greater second aquifer in upper area of flow line cross the the value of both indices reflects the greater impact of Japanese standard value such as KW-39 and KW-40 cation exchange in the groundwater (Li et al. 2015). The (also see Supplementary data). However, it can be seen indices of groundwater vary from -9.94 to 0.13 and from Fig. 1 that majority of the high-fluoride wells are -1.02 to 0.48 for CAI 1 and CAI 2, respectively. situated at the stagnant area, i.e., Kumamoto Plain area, Figure 3 shows that almost all the deep and shallow where about 57 % of the collected samples that contain more than Japanese standard limit indicate that wide exchange, where Na and K+ of the aquifer are exchang- range of fluoride enrichment occurred at Kumamoto ing with Mg?+ and Ca?+ in the groundwater. Masue et al. Plain area only. Tsuru et al. 2006 also found highest (2007), McNab et al. (2009), and Li et al. (2015) also value (3.3 mg/L, average of 10-year water quality- reported similar situations in different aquifer systems monitoring data) at the same location. In general, Springer Water Air Soil Pollut (2016) 227: 385 Page 9 of 14385 highest F- concentrations are accumulated along the last 4.2 Sources and Processes Controlling F Enrichment marginal line of the plain area, and shallow wells con- in Groundwater tain higher F- concentrations than deeper wells (see Supplementary figure). Average molar ratios of F/Cl in The high-F occurrence in Kumamoto basin is narrowly shallow and deep groundwater are 0.08 and 0.06, re- distributed in the Kumamoto Plain area and associated spectively, which is nearly comparable with unpolluted with possibly high-pH, high-HCO3, and high-Na+ rainwater F /C1 ratio ~0.02 (Saether et al. 1995) and groundwater characteristics. Generally, fluoride level therefore indicates that geochemical processes are likely in the groundwater largely depends on the availability controlling the fluoride enrichment in this area. of F-bearing minerals, and their leaching from the rock formation is one of the major processes for F enrich- ment (Ayenew 2008). The main F minerals in igneous rocks are fluorite, apatite, micas (biotite), and horn- 4 Discussion blende (Jacks et al. 2005). Kumamoto volcanic basin is dominated with silicate, and alumino-silicate minerals 4.1 F- Geochemistry in Groundwater and hydrolysis of plagioclase minerals, pyroxenes, and biotite are the dominant weathering processes, which The geochemistry ofF and certain major ions as well as greatly control the groundwater chemistry in Kumamoto pH in the Kumamoto basin is very distinctive. Table 2 basin (see in study area settings). Therefore, these min- shows the average hydrochemistry of high F erals could be the primary source of dissolved F in (F >0.8 mg/L) and low F (F <0.8 mg/L). It can be groundwater even though F concentrations in ground- seen from the table that high F is related with high-Na*, water that derive through leaching process from aquifer low-Ca2+, high-HCO3, high-pH, and high-TDS matrix are generally low compare to the F concentra- groundwater. Figure 4 also supports that high-F tion in the matrix minerals. Kumamoto groundwaters groundwater is mainly Na-HCO3 type. Average Nat are alkaline in nature due to high pH; it contains rela- and TDS value of high-F groundwater are around two tively high OH , which can provide high anion ex- times higher than from low-F groundwater. Arsenic change capacity by replacing F from the matrix min- and phosphate concentrations are also much higher in erals and thus increasing F concentration in the ground- high-F groundwater than in low-F groundwater. F water. The reaction can be written as follows: shows a positive correlation with Na+ and Na/Ca ratio, in which r2 values are 0.59 and 0. 53, respectively, KMg3[AISisO10] F2 + 2OH (1) ignoring few outlier points (Fig. 4a, b). Correlations of = KMg[AISisO1o][OH]2 + 2F- + 2Mg2+ F with pH and HCO3 are relatively weak, and the coefficients are also insignificant (Fig. 5c, d). From Fig. 5, it is obvious that F mobilization in However, a stronger positive correlation is found be- groundwater is facilitated by high Nat/Ca?+ ratios. tween F and PO43- (r2 =0.92) (Fig. 5e). F and As(T) These high Na/Ca ratios are the results of cation ex- concentrations in first aquifer shows a positive correla- change, which is one of the dominant processes for tion, although this relation in second-aquifer samples is groundwater evolution at the study area especially in widely scattered (Fig. 5f). All these scatter plots and the Kumamoto Plain area due to its stagnant nature. their relations illustrate that variations of F- in the Currell et al. (2011) and Masue et al. (2007) had been groundwater inherently associated with major ion chem- conducted different experiments to investigate the effects istry and thus appeared as an important control for the of Na+ and Ca?+ on F release, and both concluded that mobilization of F higher concentration of F was released in presence of Table 2 Average hydrochemistry of high-F (>0.8 mg/L) and low-F (<0.8 mg/L) groundwater Average (T/sn) sV (T/su) sTL (T/s) tOd (T/) OOH (T/) (/bw) eN (Tbu) Hd (u) aldues High-F groundwater 17 7.981.08 119.3 11.7 176.9 4.4 466 10.86 Low-F groundwater 3 7.710.42 49.1 18.1 110.9 0.5 261 5.73 Springer 385Page 10 of 14 Water Air Soil Pollut (2016) 227: 385 Fig. 4 Piper diagram showing 100 Legend the groundwater characteristics and evolution in the study area. F >0.8 mg/L Samples containing F ○ 1st aquifer concentration above Japanese 2nd aquifer drinking water standard limit (F >0.8 mg/L) marked separately 100 Ca2+ CI CATIONS ANIONS high Nat/Ca2+ solution compared to the low Na+/Ca2+ sulfate reduction and removes from the system as arseno- solution. Li et al. (2015) described in details how high pyrite. It is evident that reduced form of As was removed Nat/Ca?+ groundwater promotes the F release into the from the groundwater as pyrites in the deeper ground- groundwater in Yungcheng Basin in northern China; it is water of western part of Kumamoto (Hossain et al. either increasing the positive surface charges of sedi- 2016), which may be the reason of lack of correlation ments or impacting in dissolution/precipitation of between As and F in second-aquifer samples. In addi- calcite and fluorite in groundwater. We assume that tion to this, poor correlation between pH and F shows similar phenomenon is prevailing in Kamamoto Basin. that pH alone does not control the fluoride enrichment On the other hand, Kim et al. (2012) strongly argued that but facilitate the process. Furthermore, presence of Na- increasing of F in reducing environment is due to the rich groundwater accelerates this process by decreasing reductive dissolution of Fe-(hydro)xides rather than positive surface density around hydrous metal oxides leaching. Our previous study (Hossain et al. 2016) re- (McNab et al. 2009; Scanlon et al. 2009). Therefore, vealed that Fe and Al oxide-hydroxides are the potential again cation exchange seems an important control that adsorbents present in this area and negative ions such as creates ambient condition for the mobilization of F in arsenate (HAsO42-) adsorb on the surface sites of these the groundwater. Another possible factor which may minerals. Moreover, relatively high pH in the groundwa- either facilitate the desorption or limit the adsorption of ter was found as one of the controlling factors in desorp- HAsO4²- and F to the metal surface sites is the presence tion of the As species from the mineral surfaces along of competing anions. HCO3 was mentioned as a com- with heterogenic redox condition. This observation may petitor to F sorption. High HCO3 contents enable high also be acceptable for F. Substantial correlation be- F into the groundwater (Su and Puls 2001). Apart from tween As(T) and F in shallow groundwater (Fig. 5f) this, PO43- is also highly competitive with arsenate indicate possibly similar desorption process for both As (HAsO4') for sorption sites in presence of other com- and F- in the groundwater, although deeper groundwater peting agents like HCOs and silica (Meng et al. 2002; showed poor correlation. In reducing environment, As Smedley et al. 2005). Since the desorption processes of species decreases with increasing aquifer depth due to HAsO42- and F are likely similar (Kim et al. 2012), thus Springer WaterAirSoilPollut(2016)227:385 Page11of14385 2.0 2.0 (a) (b) 1.6 1.6 1.2 R²=0.53 0.8 0.0 0.0+ 100 200 300 400 500 6 12 18 24 Q 30 Na (mg/L) Na/Ca(meq/L) 2.0 2.0 (c) (d) 1.6 1.6- 1.2 （/6w） (mg/L) 0.8 0.8 0.4 0.4 ? 0.0 0.0 7.0 7.5 8.0 8.5 0'6 9.5 100 150 200 50 250 300 350 pH HCO: (mg/L) 2.0 2.0 (e) ist aquifer (f) 2ndaquifer 1.6 1.6 1.2 (7/bw) 0.8 0.4 0.0 0.0 12 10 20 30 40 50 60 70 PO。(mg/L) As(m;(μg/L) Fig. 5 Relationship between F concentration and Na+ (a), Na/Ca ratio (b), pH (c), HCO (d), PO4”- (e), and F/Cl vs As/Cl (f), respectively, in the groundwater samples Springer 385Page 12 of 14 Water Air Soil Pollut (2016) 227: 385 presence of PO43- and HCO in the groundwater may 2 also facilitate desorption of F. Phosphate in the ground- water can be introduced with the dissolution of apatite, especially the fluorapatite mineral according to the fol- lowing reaction: Cas(PO4)3 F + 6CO2 + 6H2O S = 5Ca2+ + 3H2PO4 + F + 6HCO3 (2) 4 Slcalcite A number of groundwater samples are enriched with Sl dissolved PO43- from Kumamoto Plain area, and their 5- significant positive correlations with F point towards 0.0 0.5 1.0 1.5 2.0 apatite source. Furthermore, there is a possibility of F F (mg/L) accumulation onto the clay surface by replacing phos- Fig. 6 Saturation index plots of calcite and fluorite against F phates through biogeochemical process in presence of concentration in groundwater overlying marine sediments and thick clay layer in the shallow groundwater aquifer of Kumamoto Plain area. 5 Conclusions Fluoride concentrations in groundwater may also get enriched by the CaF2 dissolution. In presence of high Rock-water interactions, groundwater geochemistry, HCO3 , fluorite dissolution reaction contributes F to and hydrogeological settings of the study area are the groundwater as follows: controlling the fluoride enrichment in the groundwater of Kumamoto area. Silicate and aluminosilicate min- CaF2 + 2HCO3-= CaCO3 + 2F + H2O + CO2(3) erals (hornblende and biotite) and apatite present in Aso pyroclastic flow deposits that consists the ground- However, there is a moderate negative correlation ss rd q po e s u ne na between F and Ca2+ in groundwater samples, meaning of fluoride. Leaching or exchange of F with OH that when Ca2+ increases in the groundwater, F de- from fluoride-bearing minerals such as fluorapatite, biotite, and desorption from metal oxide surfaces are creases (data not plotted). This phenomenon can be explained by the fluorite solubility as it is the least the major F introduction processes in groundwater. soluble F mineral. The reaction can be written as Spatial distribution of fluoride shows that fluoride follows: enrichment mainly occurred in the Kumamoto Plain area, where 57 % of the collected groundwater sam- 2F- + Ca2+ = CaF2 ples exceeds Japanese drinking water permissible limit (4) for fluoride (>0.8 mg/L). Comparatively, high fluoride is accumulated in the first aquifer than in the second Equation (4) implies that ifCa2+ concentrations reach aquifer. High F is noticeably associated with high pH, in saturation with respect to fluorite, F will be de- high HCOs , high Na/Ca ratios, and Na-HCO3-type creased through mineral precipitation. However, Fig. 6 groundwater. Cation-exchange process poses the dom- shows that although initial fluorite has a positive ten- inant control on groundwater fluoride mobilization and dency to reach in equilibrium, groundwater is still un- enrichment due to decrease of positive-charge density dersaturated with respect to fluorite, which means that around metal oxide surfaces. High pH values and precipitation of fluorite does not limit the F concentra- HCO3, though alone cannot control dominantly, fa- tion in groundwater. Nonetheless, few samples were cilitate the desorption process. Presence of other com- reached in calcite saturation indicating the possibility peting anion PO4’-, which is probably due to apatite of calcite precipitation. If so, it could facilitate F en- dissolution, may magnify this desorption process. richment according to Eq. (3) as the reaction progress Calcite and fluorite are the major limiting minerals further rightward. that control the concentration of high F- in the Springer Water Air Soil Pollut (2016) 227: 385 Page 13 of14385 groundwater. Finally, stagnant character of Kumamoto Deshmukh, A. N., Valadaskar, P. M., & Malpe, D. B. (1995). Fluoride in environment: a review. Gondwana Geological Plain area which creates a particular geochemical con- Magazine, 9, 1-20. dition is a critical factor for fluoride enrichment in the Diaz-Barriga, F., Leyva, R., Quistian, J., Loyola-Rodriguez, J. B. groundwater of western Kumamoto basin. Pozos, A., & Grimaldo, M. (1997). Endemic fluorosis in San Luis Potosi, Mexico. Fluoride, 30, 219-222. Dissanayake, C. B. (1991). The fluoride problem in the ground- Acknowledgments This research was financially supported water of Sri Lankaenvironmental management and health. 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Bulletin of Nordstrom, D. K., Ball, J. W., Donahoe, R. J., & Whittemore, D. Kumamoto city environmental research institute, 52-63 (in (1989). Groundwater chemistry and water-rock interactions Japanese). at Stripa.Geochimica et Cosmochimica Acta, 53,1727- Watanabe, K. (1978). Studies on the Aso pyroclastic flow deposits 1740. in the region to the west of Aso caldera, southwest Japan, I: Ono, K., & Watanabe, K. (1985). Geological map of Aso Volcano Geology of the Aso-4 pyroclastic flow deposits. Mem. Fac. (1:50,000). In Geological map of volcanoes 4. Geological Educ., Kumamoto Univ., 27, Nat. Sci., 97-120. survey of Japan (in Japanese with English abstract). WHO. (2004).Fluoride in drinking water—background document Parkhurst, D. L., & Appelo, C. A. J. (1999). User's guide to for development of WHO guidelines for drinking water qual- PHREEQC (version 2)-a computer program for speciation, ity. 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The Islaizd Arc (1996) 5, 289-320
Thematic Article
Anatomy and genesis of a subduction-related orogen:
A new view of geotectonic subdivision and
evolution of the Japanese Islands
YUKIO ISOZAKI
Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 0-okayama,
Megur.0, Tokyo 152, Japari
Abstract The Japanese Islands represent a segment of a 450 million year old subduction-related
orogen developed along the western Pacific convergent margin. The geotectonic subdivision of
the Japanese Islands is newly revised on the basis of recent progress in the 1980s utilizing
microfossil and chronometric mapping methods for ancient accretionary complexes and their
high-P/T metamorphic equivalents. This new subdivision is based on accretion tectonics, and it
contrasts strikingly with previous schemes based on ‘geosyncline’ tectonics, continent-continent
collision-related tectonics, or terrane tectonics. Most of the geotectonic units in Japan are
composed of Late Paleozoic to Cenozoic accretionary complexes and their high-PIT metamorphic
equivalents, except for two units representing fragments of Precambrian cratons, which were
detached from mainland Asia in the Tertiary. These ancient accretionary complexes are identified
using the method of oceanic plate stratigraphy. The Japanese Islands are comprised of
12 geotectonic units, all noted in southwest Japan, five of which have along-arc equivalents in
the Ryukyus. Northeast Japan has nine of these 12 geotectonic units, and East Hokkaido
has three of these units. Recent field observations have shown that most of the primary
geotectonic boundaries are demarcated by low-angle faults, and sometimes modified by second-
ary vertical normal and/or strike-slip faults. On the basis of these new observations, the tectonic
evolution of the Japanese Islands is summarized in the following stages: (i) birth at a rifted
Yangtze continental margin at ca 750-700 Ma; (ii) tectonic inversion from passive margin to
active margin around 500 Ma; (iii) successive oceanic subduction beginning at 450 Ma and
continuing to the present time; and (iv) isolation from mainland Asia by back-arc spreading at ca
20 Ma. In addition, a continent-continent collision occurred between the Yangtze and Sino-
Korean cratons at 250 Ma during stage three. Five characteristic features of the 450 Ma
subduction-related orogen are newly recognized here: (i) step-wise (not steady-state) growth of
ancient accretionary complexes; (ii) subhorizontal piled nappe structure; (iii) tectonically
downward-younging polarity; (iv) intermittent exhumation of high-P/T metamorphosed accre-
tionary complex; and (v) microplate-induced modification. These features suggest that the
subduction-related orogenic growth in Japan resulted from highly episodic processes. The
episodic exhumation of high-P/T units and the formation of associated granitic batholith (i.e.
formation of paired metamorphic belts) occurred approximately every 100 million years, and the
timing of such orogenic culmination apparently coincides with episodic ridge subduction beneath
Asia.
Key words: accretionary complex, Japan, microfossil mapping, microplate modification, oceanic
plate stratigraphy, orogeny, paired metamorphic belts, ridge subduction, subhorizontal nappe,
Yangtze.
Accepted for publication April 1996
290 Y. Isoxaki
INTRODUCTION
The modern Japanese Islands geographically com-
prise five island arcs: the Kurile, Northeast Japan,
Izu-Bonin, Southwest Japan, and Ryukyu arcs,
showing complex patterns common in the western
Pacific. Four of these form segments of active island
arcs between the Eurasian continent and the Pacific
Ocean where the seafloor is currently subducting
westward beneath Asia; the Izu-Bonin arc is the
exception and forms an intra-oceanic arc (Fig. 1).
Ongoing subduction processes along these margins
add materials to and modify tectonic features of the
Asian continental margin. Geological studies in the
1980s revealed that the Japanese Islands evolved
under similar tectonic processes to those active
today, and that major geologic units exposed on the
islands are subduction-related products of the Late
Paleozoic to Cenozoic orogenies.
It appears that the orogenic belts in Japan have
widened oceanward by about 400 km in 450 million
years, by virtue of long-term subduction of the
Pacific seafloor along the Yangtze (South China) and
Sino-Korean (North China) continental blocks. Ac-
cording to the classic categorization of orogenic
belts by Dewey and Bird (1970), this 450 million
year old orogen of Japan corresponds to a typical
example of a Cordilleran-type orogen between a
converging pair of continental and oceanic plates,
that features a subduction complex, high-P/T
schists and coeval granitic batholiths. The Cordil-
leran-type orogen is thought to originate from
steady-state subduction of an oceanic plate beneath
a continental plate, and contrast was emphasized
between the Cordilleran-type and the Alpine-Hima-
layan style (or continent-continent collision-type)
orogen. Although great variation in orogenic styles
and possible plate tectonic mechanisms within the
so-called Cordilleran orogens has since been de-
scribed, the distinction between the Cordilleran-type
and collision-type orogens still appears justified.
Geologic and tectonic studies in Japan during the
last two decades have clarified several new and
significant aspects of oceanic subduction-related
(or classic Cordilleran-type) orogens. New observa-
tions are highlighted in the following five charac-
teristic tectonic features of Japan: (i) intermittent
growth of accretionary complexes; (ii) subhorizon-
tal nappe structure; (iii) downward younging polar-
ity; (iv) episodic formation of a tectonic sandwich
with a high-PIT unit; and (v) microplate modifica-
tion. Particularly noteworthy is the episodic (non-
steady-state) growth pattern of a subduction-
related orogen without a collision of continent or arc because it clearly contrasts with the previous
understandings on steady-state growth of the
Cordilleran-type orogen.
This article reviews the latest version of the
geotectonic subdivision of the Japanese Islands in
view of these developments and discusses its tec-
tonic implications. The implications of this review
appear to impact also on subduction-related oro-
gens in general. The new geotectonic subdivision is
fundamentally adopted from the summary by Iso-
zaki and Itaya (1991) and Isozaki and Maruyama
(1991) published in Japanese, and is slightly mod-
ified to accommodate recent information. A histor-
ical review of studies of orogeny and geotectonic
subdivision in Japan is also given in the appendix.
OCEANIC PLATE STRATIGRAPHY (OPS) ANALYSIS FOR
ANCIENT ACCRETIONARY COMPLEX (AC)
Nearly 90% of the shallow-level crust of the Japa-
nese Islands is occupied by Late Paleozoic to Meso-
zoic accretionary complexes and granitic batho-
liths. This observation suggests that the major
orogenic framework of the islands formed at this
time, and that the geotectonic subdivision of the
Japanese supracrust depends mainly on the 3D
configuration of the accretionary complexes and
their high-PIT metamorphosed equivalents.
In this article, the term accretionary complex (AC)
is strictly used in the following sense: an AC is a
geologic entity which grows in situ in trench and
trench inner wall in an active subduction zone as a
result of subduction-driven layer-parallel shortening
and vertical stacking/thickening of trench-fill mate-
rials usually composed of oceanic sediments and un-
derlying volcanic rocks. This definition excludes con-
tinental blocks or island arcs, even though they once
occurred in an oceanic domain prior to arrival at the
continental margin. The term OPS is an acronym of
‘oceanic plate stratigraphy’ and this represents a
sequence of sediments and volcanic rocks accumu-
lated primarily on an oceanic plate prior to sub-
duction-accretion at trench (Fig. 2). A full spectrum
of an ideal OPS is comprised of, in ascending order:
(i) MORB basalt; (ii) pelagic/hemipelagic sediments;
and (iii) trench-fill turbidites, similar to the rocks
recovered from drilling through modern trench
floors. In modern examples, almost identical OPS is
often preserved also in imbricated thrust packages
in the trench inner-wall, that is, the youngest part in
an AC. This unique stratigraphy was once called
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Anatomy and evolution of Japanese Islands 291
/II \ N. American Piate
Fig. 1 Plate tectonic framework around the
present day Japanese Islands Major plate in-
teractions are oceanic subductions along the
Nankai trough off southwest Japan and along
the Japan trench off northeast Japan The
former is an accretionary margin while the
latter is an erosional margin with a high con-
vergence rate and rugged-surfaced oceanic
plate The Mariana trench represents another
non-accretionary margin According to the ob-
lique subduction of the Pacific and Philippine
Sea plates southwestward moving fore-arc
slivers develop in 3 domains adding second-
ary across-arc structural features The Miocene
back-arc basin Japan Sea, is under destruc-
tion along the convergent plate boundary be-
tween Eurasian and North American plates
while a new back-arc basin is emerging in the
Ryukyus associated with rifting of the iconti-
nental crust
‘plate stratigraphy’ (Berger & Winterer 1974), how-
ever, the adjective ‘oceanic’ is added later in order to
exclude sediments derived and accumulated exclu-
sively on continental plate. Details of the concept of
OPS and the practical example of the OPS analysis
can be found in Isozaki et al. (in press b) and Mat-
suda and Isozaki (1991).
It is generally easy to recognize the occurrence of
modern AC by seismic resea,rch simply because they
occur immediately next to active trenches. Ancient
examples exposed on land, on the contrary, have
usually lost contact with their primary tectonic set-
ting nor wedge-like 3D geometry through later over-
printing processes, thus their recognition requires more specific information, not by external geometry
but by internal structure and composition. The most
reliable way to identify ancient AC, and to distin-
guish them from neighboring ones, is the ‘OPS an-
alysis’ refined mainly in Japan in the 1980s.
Ancient AC exposed on land also possess OPS
which usually consists of a sequence in ascending
order of greenstones (less than several tens of
metres thick), deep-sea pelagic chert (less than
200 m thick), hemipelagic siliceous claystone (less
than 100 m thick), and terrigenous clastic rocks
such as mudstone and sandstone (more than 200 m
thick). As most AC formed by the same teconic pro-
cess from deep-sea rocks and sediments, they are
rands1 lmudst
IC~OUS mudstone
bedded chert - birth at MOR demlae at
I s~bducllon zone t Fig. 2 Simplified ridge subduction sys-
tem and the concept of Oceanic Plate
stratigraphy (OPS) (modified from Mat-
suda & lsozaki 1991) Note the age gap
between two distinct horizons (solid tri-
angle the horizon between pillowed MORB
greenstone and pelagic chert marking the
birth of oceanic plate open triangle the
and terrigenous clastics marking the arrivdl
at trench) which represents the total travol
time of the subducting oceanic plate from
mid-oceanic ridge to trench in othu
words the age of the subducting oceanic
plate at trench auctlon zone
horizon between hemipelagic mudstone Wedge mantle
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
292 Y. Isoxaki
best distinguished by accurate age determination.
Documenting a well-dated OPS for an ancient AC
provides significant information on the timing of ac-
cretion and the age of the subducted ancient oceanic
plate. Microfossil (conodont and radiolaria) dating of
component sedimentary rocks is particularly effec-
tive in OPS analysis for ancient AC exposed on land
because it reveals (i) the accretion age at trench
( = age of the horizon between hemipelagic sediment
and trench-fill turbidite); and (ii) an age of sub-
ducted oceanic plate responsible for the accretion
( = duration of pelagic + hemipelagic deposition)
(Fig. 2). Thus if documentation of a unique OPS is
available with precise dating, it can provide a prime
identity for each ancient AC unit that appears as a
look-alike to neighboring units.
The OPS analysis for on-land exposed ancient AC
has played the main role in the drastic change in
geological studies in Japan in the 1980s (Ichikawa et
al. 1990), and the research style of detailed field
mapping in 1:5000 scale combined with microfossil
dating is here called ‘microfossil mapping’ of ancient
AC. An interval of one microfossil zone is less than 5
million years in average for the late Paleozoic to
Mesozoic, although their resolution for dating is less
precisely controlled with few tie points in time scale.
However, it is still generally quite useful to distin-
guish neighboring units in the field.
In addition, ages of metamorphosed AC of the
‘grey zone’ (see Appendix) that make microfossil
dating difficult are determined radiometrically using
various methods (K-Ar, Ar-Ar, Rb-Sr, and fission
track). In some cases, an AC is accurately dated by
two ‘radio ages’ (i.e. radiolarian-based accretion age
and radiometrically dated age of subduction-related
regional metamorphism). The subduction-related re-
gional low-grade metamorphism usually occurred
-10-20 million years later than the former (Takami
et al. 1990, 1993; Kawato et al. 1991; Isozaki in
press b), and it can add another reference feature for
comparison in OPS. This research style combining
geochronology dating and detailed field mapping is
here called ‘chronometric mapping’ of ancient AC.
Practical examples of these research styles ap-
plied to the Permian and Jurassic AC in SW Japan
are reviewed in detail in Isozaki (in press a,b) and
Kimura (1996).
GEOTECTONIC SUBDIVISION
The newly proposed geotectonic subdivision of the
Japanese Islands based on these methods (Figs 3,4)
is described here in four sections, based on domains in southwest Japan, the Ryukyus, northeast Japan,
and east Hokkaido that are at present separated
physiographically and/or tectonically by secondary
transverse faults. The fundamental scheme of this
subdivision is after Isozaki and Maruyama (1991),
and is slightly modified according to the latest infor-
mation. The Izu-Bonin Islands are not described in
this article, as they form a young island arc of intra-
oceanic nature and have lesser significance to the
primary orogenic framework along the Asian conti-
nental margin.
In southwest Japan, including Kyushu, Shikoku
and western Honshu Islands, the Cenozoic volcano-
sedimentary covers are thinner than those in other
domains due to the rapid uplift in the Quaternary,
and this allows extensive exposure of Late Paleo-
zoic to Mesozoic AC and their metamorphic equi-
valents. The Ryukyus and northeast Japan can be
essentially treated as lateral extension of south-
west Japan, however, these domains were consid-
erably modified and dislocated by secondary tecton-
ism including movement of fore-arc sliver, back-arc
spreading, and arc-arc collision. Thus these two
domains will be briefly explained after southwest
Japan as its lateral equivalents. The island of
Hokkaido is characterized by a unique setting with
an arc-arc collision between the Northeast Japan
arc and Kurile arc. East Hokkaido that belongs to
Kurile arc will be mentioned separately.
In the course of explaining geotectonic subdivi-
sion, the term ‘belt’ is used in this article to describe
the distribution of geotectonic units in two dimen-
sions. When a three-dimensional geologic entity is to
be described, non-genetic terms like unit, body,
block, complex, and nappe are used. The term ‘ter-
rane’ is avoided here because it may be confused
with a prejudicial connotation of allochthoneity (Co-
ney et al. 1980; Howell 1985), which has been sug-
gested by Sengor and Dewey (1991) and Hamilton
(1991). Refer to Ichikawa et al. (1990) for more
detailed descriptions and relevant references of pre-
Cretaceous belts; and to Taira et al. (1988,1989) for
those of much younger belts. In this long description
section, readers interested in tectonics rather than
local geology of the islands may read only the de-
scription of southwest Japan and skip those of the
Ryukyus, northeast Japan, and East Hokkaido be-
cause the latter areas represent lateral equivalents
of southwest Japan.
SOUTHWEST JAPAN
In southwest Japan the two-dimensional east-west
trending zonal arrangement of various geologic
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Anatomy and evolution of Japanese Islands 293
Fig. 3 New geotectonic subdivision of the Japanese islands (modified from lsozaki & Maruyama 1991). The geotectonic subdivision of the Ryukyus and their
correlation are shown separately in Fig. 8. Explanatory symbols for geotectonic units used in the figure and text are as follows: Southwest Japan [Rn. Renge belt ~
Sn: Sangun b ; Ak: Akiyoshi b , Mz. Maizuru b ; UT- Ultra-Tanba b (included in Mz in text), M-T. Mino-Tanba b., Ry. Ryoke b ; Sb: Sanbagawa b ; Ch: Chtchibu
b (including Northern Chichibu b.; Kurosegawa belt, and Southern Chichibu belt in text); Sh; Shimanto b. (divided into Northern and Southern Shimanto belts in
text]; Northeast Japan [Ht-Tk: Hitachi-Takanuki b. ( = Hida b.); Gs. Gosaisho b. ( = Ryoke b.), MM: Matsugataira-Motai b. ( = Renge b ); SK. Southern Kitakamt
b. ( = Oki b.); NK. Northern Kitakami-Oshima b. ( = Mino-Tanba b.); Kk: Kamuikotan b. ( = Sanbagawa b.), Hdk. Hidaka b. ( = Shimanto b.), Tokoro b
( = Sanbagawa b. + Shimanto b.); Nm: Nemuro b.]. Solid black area represents ophiolitic zone
units is more obvious if surface cover and granitic
intrusions are removed. Southwest Japan com-
prises 12 distinct geotectonic units (Fig. 5); from
oldest to youngest; a 2.0 (;a-250 Ma gneiss com-
plex; a 230 Ma intermediate-pressure type meta-
morphic complex; a 580-450 Ma ophiolite; a 400-
300 Ma high-P/T schist; a 250 Ma AC; a 230-
200 Ma high-P/T schist; a 180-140 Ma AC; a
120-100 Ma low-P/T metamorphic complex; a
100 Ma high-P/T schist, an 80 Ma AC; and a
40-20 Ma AC. These units are distributed in 15
belts, that is, from the Japa.n Sea side to the Pacific
side: Oki belt; Hida b.; 0-eyama b.; Renge b.;
Akiyoshi b.; Sangun b.; Maizuru ( + Ultra-Tanba)
b.; Mino-Tanba b.; Ryoke b.; Sanbagawa b.;
Northern Chichibu b.; Kurosegawa b.; Southern
Chichibu b.; Northern Shimanto b.; and Southern
Shimanto b. Repeated occurrence of the same unit
in more than two belts (i.e. outliers in the form of klippes and/or tectonic windows as shown in
Fig. 6) in southwest Japan causes a mismatch in
the number of belts and geotectonic units. For
example, the Jurassic AC apparently occur in three
belts, that is, the Mino-Tanba b.; N. Chichibu b.;
and Southern Chichibu b., separated from each
other for up to 50 km, although these are identical
in terms of OPS.
The three belts along the Japan Sea side, that is,
the Oki, Hida and 0-eyama belts, are intimately
linked to Precambrian crusts in nature, and they
form the core of the Phanerozoic orogen in south-
west Japan (Fig. 3). On the other hand, the other
12 belts surrounding the above three represent
zones of subduction-related accretionary growth
that account for the 450 million year-long widen-
ing and thickening of southwest Japan.
The Oki belt of continental affinity is at present
isolated from mainland Asia, as it was rifted and
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
294 Y. Isoxaki
A A’ Hokkaido
B’
/ /
C’
h / /
/
S Ryukyus
D D’
\fib Pacific Ocean
C’
/ / / NE Japan
SW Japan
1 lOkm
100km
Granites a Cretaceous and
younger ACs
J Jurassic ACs
P Pre-Jurassic ACs
Continent
Shelf sediments
Fig. 4 Across-arc geologic profiles of the Japanese islands along four transects Hokkaido NE Japan SW Japan and southern Ryukyus (modified from lsozaki &
Maruyama 1991) Abbreviations for major boundary faults are as follows (I) Thrust (bold) HMT Hidaka Main Thrust HyTZ Hayachine tectonic zone Ng-Hm TL
Nagato-Hida marginal tectonic line, I-KTL Ishigaki-Kuga tectonic line BTL Butsuzo tectonic line (11) Strike-slip faults ATL Abashiri tectonic line, TTL Tanakura
tectonic line HTL Hatagawa tectonic line H-KTL Hizurne-Kesen nurna tectonic line, MTL Median tectonic line
detached by the Miocene opening of a back-arc
basin, the Japan Sea (Jolivert et al. 1994; Otofuji
1996, this issue). Judging from the lithologic/
chronologic similarity to the Precambrian rocks of
the Sobaesan massif in South Korea, the Oki belt is
regarded as an eastern extension of the Yangtze
(South China) craton (Sohma et al. 1990; Isozaki
& Maruyama 1991) which represents one of the
continental pieces rifted apart from the supercon-
tinent Rodinia at 750-700 Ma (Powell et al. 1993;
Li et al. 1995).
On the other hand, the kyanite-bearing Hida
metamorphic rocks are regarded as the northeast-
ern extension of the ultrahigh-pressure to high-
pressure metamorphic rocks along the Qinling-
Dabie suture (230 Ma continent-continent collision
zone) in central China between the Yangtze and
Sino-Korean (North China) blocks (Wang et al. 1989; Maruyama et al. 1994; Cong & Wang 1995).
The unique occurrence of Middle to Late Paleozoic
shelf strata with Boreal fauna in the periphery of
the Hida belt (Igo 1990; Kato 1990) suggests a
strong link between the Hida belt and Sino-Korean
block. The present position of the Hida belt in the
central part of Japan is due to later across-arc
contraction and juxtaposition (Komatsu 1990), and
the primary contact with the Oki belt (the Yangtze
block) has been lost. Concerning the geotectonic
correlation of the Oki and Hida belts with continen-
tal blocks and later tectonic juxtaposition, refer to
Isozaki (in press a).
The 450-580 Ma ophiolite of the 0-eyama belt
along the southern margin of the Oki belt is the
oldest oceanic material in Japan. As its easterly
extension in northeast Japan has mid-Paleozoic
sedimentary cover of continental shelf facies
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Anatomy and evolution of Japanese Islands 295
Fig. 5 Geotectonic subdivision of southwest Japan (modified from lsozaki & ltaya 1991) Hatched areas represent occurrence of Paleozoic AC, and solid black
areas indicate their tectonic outliers on the Pacific side ( = Kurosegawa b) The Nagato-Hida Marginal tectonic line (b) separates the continental Oki belt from the
younger Paleozoic-Mesozoic AC belts The lshigaki-Kuga tectonic line (f) divides Paleozoic AC belts and Mesozoic ones in southwest Japan Note its winding
surface trajectory indicating a low-angle nature as shown in Fig 6
(Ozawa 1988), the 0-eyama ophiolite probably
represents a remnant of the initial oceanic crust
that formed through the break-up of Rodinia at ca
750-700 Ma. Prior to the subduction regime
started from 450 Ma at the latest, a piece of the
initial Pacific ocean floor was likely attached to the
rifted continental margin of the Yangtze block;
that is, the Oki belt (Isoza,ki & Maruyama 1991).
The occurrence of the medium-pressure-type am-
phibolite in the 0-eyama belt probably suggests its
involvement in the 250 Ma collision event (Isozaki
in press a). The lithologic assemblage, age, and
other characteristics of these three units, with
continental affinity, are briefly described.
1, Hida belt (Hd); 1.1 Ga to 250 Ma medium-
pressure-type metamorphic rocks and 180 Ma
granites (Sohma & Kunugiza 1993). The highest
metamorphic grade reaches the upper amphibolite
facies, only locally to the granulite facies. Proto-
liths of paragneiss are composed of sedimentary
rocks that most likely accumulated along the pas-
sive continental margin. These include peralumi-
nous siliciclastic rocks and impure carbonates with
a minor amount of mafic igneous rocks. Non- to
weakly metamorphosed Middle to Late Paleozoic
shelf sequences occur fragmentally in the periphery
of the belt; 2, Oki belt (Ok): 2.0 Ga to 250 Ma
medium-pressure-type gneiss and granite complex
(Suzuki & Adachi 1994). The highest metamorphic
grade reaches the upper amphibolite to granulite
facies. Protolith is continental sedimentam rocks with minor amount of mafic igneous rocks; 3,
0-eyama belt (Oe): 450 to 580 Ma dismembered
ophiolite composed mostly of serpentinized ultra-
mafic rocks (lherzolitic.harzburgite), metagabbros,
and a fragment of medium-pressure-type amphib-
olite (Arai 1980; Kurokawa 1985).
The belts 4-15 listed below are zones of
subduction-related accretionary growth that prac-
tically account for the 450 million year-long oro-
genic widening and thickening of southwest Japan
(Figs 5,6) together with underlying granitic batho-
liths emplaced later. These AC belts including
metamorphosed equivalents were successively
added to the southeastern margin of the continent,
in particular around the Yangtze block. A summary
on OPS for these AC units in southwest Japan is
shown in Fig. 7. It is noteworthy that the timing of
accretion is generally getting younger oceanward
from the 400 Ma meta-AC to the Miocene AC, and
that the age of the subducted oceanic plate respon-
sible for accretion has been considerably variable;
from -160 million years old for the Jurassic AC to
almost zero for the Southern Shimanto AC. For a
more detailed description of belts 5-7 and 12, refer
to Isozaki (1996a); for a description of belts 8, 11
and 13 refer to Isozaki (in press b). 4, Renge belt
(Rn): 400-300 Ma high-P/T schists and associated
serpentinite (Nishimura 1990). The highest meta-
morphic grade reaches the high-pressure amphibo-
lite facies through glaucophane schist facies. The
protolith is an AC of unknown age composed of
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296 Y. Isoxnki
,- Pre-Jurassic complexes ~
of the Inner Zone 12. Ks presentj
Tectonic Outlier
Naoato- Hida
N S
1.
LA INNER ZONE OUTER ZONE -- - -.
B
Fig. 6 Geotectonic profile of southwest Japan (modified from A lsozaki & ltaya 1991, B lsozaki eta/ 1992) Profile A shows the subhorizontal nature of the piled
nappes of AC with klippes and tectonic windows on the surface Note the downward younging polarity among the AC Block diagram B is an enlargement of central
portion of profile A showing correlation of geotectonic units and boundaries across the Median Tectonic Line (M T L ) Concerning the two-fold nature of the M T L
refer to text and Figs 17b c and 18b
greenstones, siliciclastics and chert; 5, Akiyoshi
belt (Ak): Late Permian (250 Ma) AC composed of
oceanic greenstones mostly of oceanic island basalt
(OIB) origin, chert, reef limestone, and terrigenous
clastics (Kanmera et al. 1990). This unit suffered
from low grade regional metamorphism up to the
lower greenschist facies at 220 Ma; 6, Sangun belt
(Sn): 230-210 Ma high-PIT schists (Nishimura
1990). The highest metamorphic grade reaches the
high-pressure amphibolite through the glau-
cophane schist facies. The protolith is an AC,
probably containing a part of the 250 Ma Akiyoshi
AC. Neighboring schist unit with problematic 200-
180 Ma ages (probably secondarily annealed) are
also included here; 7, Maizuru belt (Mz): Middle-
Late Permian AC with 280 Ma ophiolite (Hayasaka
1990; Ishiwatari et al. 1990). The unit sometimes
discriminated as the ‘Ultra-Tanba belt’ (Caridroit
et al. 1981; Ishiga 1990) is included here. The
ophiolite suite is dismembered but its primary thickness is estimated to be -25 km; 8, Mino-
Tanba belt (MT): Jurassic AC with a minor amount
of latest Triassic and earliest Cretaceous parts
(200-140 Ma) (Wakita 1988; Nakae 1993). This
unit is composed of oceanic greenstones of OIB
origin, deep-sea pelagic chert, reef limestone, and
terrigenous clastics. Secondarily mixed AC (olis-
tostromes and melanges) occur commonly. This AC
unit is tentatively subdivided into three parts; (i)
Early Jurassic part (accreted at 200 Ma; metamor-
phosed at 170 Ma); (ii) Middle Jurassic part (ac-
creted at 170 Ma; metamorphosed at 140 Ma); and
(iii) Late Jurassic part (accreted at 140-150 Ma;
metamorphosed at 120 Ma); 9, Ryoke belt (Ry):
120-100 Ma low-PIT metamorphic rocks and asso-
ciated granites (Nakajima in press). The highest
grade part includes sillimanite-bearing gneiss. Pro-
tolith is mostly composed of Jurassic AC with
lesser amount of the pre-Jurassic AC and sedi-
ments. Age of granites ranges mostly in 120-
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Anatomy and evolution of Japanese Islands 297
UbdUCted
ceanic Plate
4- !45 Fa ra /lo n
Akiyoshi
l0w.P mel. bell
100 Ma- __ Sangun
hrgh PIT
me1 belt
0 0 ,-
0 0 N
A
A
f
-300 M: lzanagi
Mino-Tanba-Chichibu (11. 13
Iow-P mel. belt
k
400 Ma-- ~__ hanba- gawa
met. bell high-PIT
~.~ 0 e?
-0 ~ ..N a:
I,
----- 4 ._
lllllllllll
llllllllll! - Kula /Pacific /Philip.Sea
& ;I 3
Northern Shimanto ~ Southern Shimanto
low-P met bell ti
metamorphism ~ Legend
* peak Of
trench-1111 turbidites
hemipelagic sediments
pelagic sediments I
I A metagabbro I
Fig. 7 Oceanic plate stratigraphy (OPS) of AC units in southwest Japan and age of subduction-related regional metamorphism (modified from lsozaki &
Maruyama 1991) Note the oceanward and tectonically downward younging polarity not only in accretion timing (age of the horizon between hemipelagic mudstone
and terrigenous clastics) but also in high-P-T and relevant low-grade metamorphism (shown by asterisk)
70 Ma with eastward younging polarity along the
Southwest Japan arc; 10. Sanbagawa belt (Sb):
high-PIT metamorphosed Early Cretaceous AC
(Banno & Sakai 1989; Takasu & Dallmeyer 1990),
well known as the Sanbagawa schists. The highest
grade reaches the high-pressure amphibolite facies
and radiometric ages concentrate in 100-80 Ma.
The high-PIT Sanbagawa belt and the low-P/T
Ryoke belt (9) form paired metamorphic belts
(Miyashiro 1961); 11, Northern Chichibu belt (Cn):
Latest Triassic to Middle Jurassic AC equivalent to
the older part of the Jurassic AC in the Mino-
Tanba belt (8) (Hada & Kurimoto 1990). This unit
is regarded as forming a tectonic outlier of the
Jurassic complex of the Mino-Tanba belt; 12,
Kurosegawa belt (Kr): Fault-bounded mixture of
the pre-Jurassic elements (Yoshikura et al. 1990;
Isozaki et al. 1992). Components of above-
mentioned belts 3 to 7 occur chaotically as slivers,
lenses and/or blocks of various sizes and shapes,
enveloped within serpentiriite matrix. As a whole,
this unit represents a tectonic outlier of the pre-
Jurassic rocks, which occurred on the Asian conti-
nent side; 13, Southern Chichibu belt (Cs): Early
Jurassic to earliest Cretaceous AC equivalent.to the younger part of the Jurassic AC in the Mino-Tanba
belt (8) and partly to that in the Northern Chichibu
belt (11) (Matsuoka 1992); 14, Northern Shimanto
belt (Shn): Late Cretaceous scarcely metamor-
phosed AC composed mostly of terrigenous elastics
with lesser amount of oceanic rocks (Taira et al.
1988). Sporadically intervened are thin tectonic
slices of melanges that include oceanic greenstones
and bedded chert within scaly argillaceous matri-
ces; 15, Southern Shimanto belt (Shs): Paleogene
and Miocene little metamorphosed AC composed
mostly of terrigenous elastic rocks (Taira et al.
1988). A minor amount of tectonic melanges occur
in this belt.
Major geotectonic boundaries in southwest Japan
(Figs 5,6) are listed with their nature. Also noted in
parentheses are well-documented examples of these
boundary faults: (a) boundary between belts 1 and 2
(Hida b./Oki b.): low-angle thrust? (Unazuki suture)
activated probably in late Mesozoic; (b) boundary
between belts 2 and 3 (Oki b./O-eyama b.), and
between belts 2 and 4 (Oki b./Renge b.): low-angle
thrust (Nagato-Hida marginal tectonic line);
(c) boundary between belts 4 and 5 (Renge b./
Akiyoshi b.): low-angle thrust (Toyogadake thrust;
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298 Y. Isoxaki
Kabashima et al. 1994); (d) boundary between belts
5 and 6 (Akiyoshi b./Sangun b.): low-angle thrust
(Kitayama thrust), (e) boundary between belts 6 and
7 (Sangun b./Maizuru b.): unexamined; (f) boun-
dary between belts 7 and 8 (Maizuru b./Mino-Tanba
b.): low-angle thrust (Ishigaki-Kuga tectonic line);
(g) boundary between belts 8 and 9 (Mino-Tanba
b./Ryoke b.): high-angle normal fault (Arima-
Takatsuki line, Iwakuni fault); essential contact:
gradual metamorphic aureole; (h) boundary between
belts 9 and 10 (Ryoke b./Sanbagawa b.): primary
low-angle thrust (Paleo-Median Tectonic Line acti-
vated in the Tertiary) and secondary high-angle
strike-slip fault (Neo-MTL active in Quaternary). (i)
boundary between belts 10 and 11 (Sanbagawa
b./N. Chichibu b.): low-angle thrust (Sasagatani
fault; Kawato et al. 1991); 0') boundary between
belts 11 and 12 (Northern Chichibu b./Kurosegawa
b.): low-angle thrust (Agekura thrust, Nakatsu
thrust) = oceanward extension of the boundary fault
f (Ishigaki-Kuga tectonic line); (k) boundary be-
tween belts 12 and 13 (Kurosegawa b./Southern
Chichibu b.): low-angle thrust (Kanbaradani
thrust) = oceanward extension of the boundary fault
f (Ishigaki-Kuga tectonic line). (1) boundary be-
tween belts 13 and 14 (S. Chichibu b./N. Shimanto
b.): low-angle thrust (Butsuzo Tectonic Line, Tsub-
uro thrust; Sasaki & Isozaki 1992). (m) boundary
between belts 14 and 15 (N. Shimanto b./S. Shi-
manto b.): low-angle thrust (Nobeoka thrust, Aki
tectonic line).
THERYUKYUS
On the basis of strong similarity of components,
the Ryukyu Islands are basically regarded as the
southwestern extension of southwest Japan
(Fig. 8), however, a north-south trending trans-
verse fault clearly separates the Ryukyus from
Southwest Japan. This fault in west Kyushu Island
sharply cuts off the zonal arrangement of south-
west Japan. Right-lateral off-set of the Late Cre-
taceous high-P/T unit in mid- to west Kyushu
suggests the boundary fault between southwest
Japan and the Ryukyus belongs to the right-lateral
strike-slip fault system as well as the Tsushima
fault and Yangshan fault in southeast Korea (Yoon
& Chough 1995), probably formed in relation to
the Miocene opening of the Japan Sea along its
western margin.
Geological information on the Ryukyu Islands is
limited by the lack of exposures but some geotec-
tonic units are well correlated with those in south-
west Japan. Most of the units in this domain are composed of Late Paleozoic and Mesozoic AC and
their metamorphic equivalents. There is no geotec-
tonic unit with continental affinity in this domain
except Early Paleozoic ophiolite in west Kyushu. Ac-
cording to dredging data, the East China Sea is un-
derlain by Precambrian rocks that probably belong
to the Yangtze (South China) craton.
The geotectonic units hitherto known from
Ryukyus are listed below. For convenience, the
numbering and symbols that were used for units in
southwest Japan are also used to describe units in
the Ryukyus: 3, metagabbro from Nomo point:
450-580 Ma ophiolite ( = Oe) (Igi & Shibata 1979;
Nishimura 1990); 6, Tomuru metamorphic rocks:
220 Ma high-P/T schists ( = Sn) (Nishimura 1990);
8, Fusaki Formation: Jurassic AC ( = MT) (Isozaki
& Nishimura 1989); 10, Yuan Formation and
Takashima schists: Early Cretaceous AC and 60-
90 Ma high-P/T metamorphic equivalents ( = Sb)
(Ujiie & Hashimoto 1983); 14, 15, Kunchan Group:
Cretaceous and Paleogene AC ( = Sh) (Osozawa
1984).
The following two belt boundaries have been
examined on land in Ryukyus; (f) boundary be-
tween belts 6 and 8 (Sn/MT): low-angle thrust
(Ishigaki-Kuga tectonic line); and (1) boundary be-
tween belts 10 and 14 (SbIShn): high-angle fault
(Butsuzo tectonic line).
NORTHEAST JAPAN
Northeast Japan comprises northeastern Honshu
and the western Hokkaido Islands, and is sepa-
rated from southwest Japan by a left-lateral
strike-slip fault called Tanakura tectonic line
(T.T.L.), and from eastern Hokkaido by a north-
south trending fault in central Hokkaido, respec-
tively (Fig. 3). Northeast Japan is also character-
ized by an apparent zonal arrangement of several
geotectonic units on the surface (Fig. 9). This
domain has been intensely modified by secondary
tectonism, however, in particular by left-lateral
strike-slip faults relevant to the Miocene opening
of the Japan Sea. This faulting leads to uncertainty
in the primary geometry and nature of contact
among the component units. However, a compari-
son with southwest Japan can help in the recon-
struction of primary features of this domain. Com-
ponent units of northeast Japan are listed through
comparison with those in southwest Japan, and
short comments will be added for recent advances.
As well as the Hida, Oki and 0-eyama belts in
southwest Japan, the three belts in northeast
Japan, that is, the Hitachi-Takanuki b., Southern
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Anatomy mid evolution of Jupanese Islands 299
/ active trench / Ishigaki-Kuga Tectonic Line lshigaki island
(6) 220~. ~IT SSM.IS
(8) Jurassk AC
Fig. 8 Geotectonic map of the Ryukyus and their link to southwest Japan (modified from lsozaki & Nishimura 1989) Inset map (B) of East Asia shows potential
along-arc continuity of the Jurassic AC belt for example for -5000 km from west Philippines to Chind/Russia border Another 1000 km northerly extension IS
examined recently in the Kamtchatsuka Koryak region northeast Russia (refer to lsozaki in press b)
Kitakami b., and Miyamori-hayachine b., have
strong continental affinities, in particular to the
Yangtze craton and the collisional suture between
the Yangtze and Sino-Korean cratons. The rest are
AC units which later accreted to northeast Japan.
1, Hitachi-Takanuki belt (Ht-Tk): 250 Ma
medium-pressure metamorphics (Tagiri 1973; Hiroi
& Kishi 1989). The protoliths include Late Paleo-
zoic sedimentary rocks accumulated on the conti-
nental shelf and volcanic rocks of bimodal charac-
teristics ( = Hd). Granite-related thermal overprint
occurred regionally at 110 Ma which is correlated
to the Ryoke metamorphism in southwest Japan
(= Ry); 2, Southern Kitakami belt (SK): 440 Ma
granite and gneiss, 350 Ma and 250 Ma granites
( = Ok) with Middle to Late Paleozoic sedimentary covers characterized by marine fauna of Australian
(Gondwanan) affinity (Suzuki & Adachi 1993;
Kawamura et al. 1990; Kato 1990); 3, Miyamori-
Hayachine belt (MH): 450 Ma ophiolite ( = Oe) with
Paleozoic sedimentary covers (Ozawa 1988; Okami
& Ehiro 1988); 4, Matsugataira-Motai belt (MM):
300-400 Ma high-PIT schists ( = Rn) (Maekawa
1981); 8, Northern Kitakami-Oshima belt (NK-
0s): Jurassic AC ( = MT) + Early Cretaceous AC
(= Sb) (Minoura 1990; Okami & Ehiro 1988); 9,
Gosaisho belt (Gs): 110 Ma low-PIT metamorphic
rocks and granites ( = Ry). Protoliths include com-
ponents of the Jurassic AC (= MT) (Tagiri et al.
1993); 10, Sorachi-Yezo belt (SY): Early Creta-
ceous AC + 100 Ma high-PIT (Kamuikotan) schists
associated with (Horokanai) ophiolite ( = Sb); 14,
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300 Y. Isoxaki
Geotectonic subdivision
of Northeast Japan
Pacific Ocean
Mi no -Ta n b a b.)
Paleozoic shelf sediments
High-P/T schists
itachi -Takanuki b. Fig. 9 Geotectonic map of northeast Japan
(modified from lsozaki & Maruyama 1991)
Names in parentheses indicate correlative units
in southwest Japan Note that the primary sub-
horizontal geotectonic boundaries are cut by a
series of left-lateral strike-slip faults that were
activated by the back-arc basin (Japan Sea)
opening in the Miocene
Idonnappu belt (Id): Early Cretaceous to early Late
Cretaceous AC ( = Sh); 14, 15, Hidaka belt (Hdk):
Late Cretaceous to Paleogene AC ( = Sh) partly
metamorphosed into low- to medium-pressure type
metamorphics associated with 50 Ma migmatite-
granite.
Although several units are well correlated to their
counterparts in southwest Japan, some of the units in southwest Japan are apparently missing in north-
east Japan, such as Permian AC (5), 200 Ma high-
P/T schists (6), and 280 Ma ophiolite (7). Their
counterparts, however, may possibly be found also
in northwest Japan in future as subsurface under-
lying units bounded by blind thrusts, owing to the
subhorizontal structure mentioned below. The geo-
tectonic boundaries between these units in northeast
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Anatomy and evolution of Japanese Islands 301
(iii). The fault (viii) (Hidaka Main thrusts) repre-
sents another example of the secondary modification
upon the primary structure, and this transpressional
fault has been activated probably by ongoing west-
ward collision of the Kurile fore-arc sliver (Kimura
1985, 1996).
Due to severe secondary modifications, primary
orogenic structures in northeast Japan have not
been fully clarified, however, an analogy in rock
type, age and OPS of AC to southwest Japan sug-
gests that a similar subhorizontal piled nappe struc-
ture also predominates in northeast Japan which
may be represented by a series of subsurface blind
thrusts (Fig. 10). Japan are listed here: (i) boundary between belts 1
and 2 (Hitachi-Takanuki b.lSouthern Kitakami b.):
unknown; (ii) boundary between belts 2 and 3
(Southern Kitakami b./Miyamori-Hayachine b.):
vertical fault but primarily low-angle thrust (west-
ern margin of the Hayachine tectonic zone); (iii)
boundary between belts 3 and 8 (Miyamori-
Hayachine b./Northern Kitakami b.): vertical fault
but primarily low-angle thrust (eastern margin of
the Hayachine tectonic zone; Tazawa 1988).
The boundaries (ii) and (iii) appear almost vertical
in outcrop but the large-scale sinuous trajectory on
the surface suggests a potentially low-angle nature
in deeper levels (Fig. 10). The zone between these
two faults has been traditionally called the Haya-
chine tectonic zone because serpentinized ophiolite
occurs in an apparently narrow belt. The eastern
margin of this zone ( = b) corresponds to the Ishi-
gaki-Kuga tectonic line (f) in southwest Japan.
(iv) boundary between belts 1 and 9 (Hitachi-
Takanuki b./Gosaisho b.): east-dipping low-angle
fault probably activated before the intrusion of the
Cretaceous granite; (v) boundary between belts 9
and 4 (Gosaisho b./Matsugataira-Motai b.): left-
lateral strike-slip fault (Hatagawa tectonic line);
(vi) boundary between belts 4 and 8 (Matsugataira-
Motai b./Northern Kitakami-Oshima b.): left-
lateral strike-slip fault (Futaba fault); (vii) bound-
ary between belts 8 and 10 (Northern Kitakami-
Oshima b./Sorachi-Yezo b ): unknown (concealed
beneath Quaternary sediments); (viii) boundary be-
tween belts 14 and 15 (Idonappu b./Hidaka b.):
east-dipping high-angle thrust (Hidaka Main
Thrust) on the surface that translates into a low-
angle one in deeper level (Ikawa et al. 1995).
The faults (v) and (vi) are typical examples of
sinistral strike-slip faults as well as the Hizume-
Kesen’numa tectonic line that activated during the
Miocene opening event of the Japan Sea along its
eastern margin. These faults clearly cut the primary
sinuous boundary faults such as the faults (ii) and
Fig. 10 Profile of northeast Japan Note
the left-lateral strike slip faults that cut pri-
mary subhorizontal piled nappe structure
that includes inferred blind thrusts EAST HOKKAIDO
East Hokkaido has a rather complicated tectonic
history compared to the rest of the Japanese
Islands, probably reflecting its peculiar geotectonic
condition, sandwiched between two Cenozoic back-
arc basins, that is, the Japan Sea and Kurile basin
(Fig. 1). The domain boundary between northeast
Japan including West Hokkaido and East Hokkaido
is inferred in central Hokkaido in the name of the
Tokoro fault or Shibetu fault, however, its precise
position, geometry and nature are unknown owing
to thick Quaternary covers in between. Three geo-
tectonic units are recognized in East Hokkaido, as
described below (Fig. 11). There is no Paleozoic
and Early Mesozoic unit in East Hokkaido, and this
implies a unique geohistory for East Hokkaido: 13,
Yubetsu belt (Yb): Cretaceous AC; 10, 14, Tokoro
belt (Tk): Cretaceous AC and high-P/T metamor-
phic equivalents; 14, 15, Nemuro belt (Nm):
Cretaceous-Tertiary shelf sequences with unknown
basement probably composed of Mesozoic-
Cenozoic crystalline rocks of arc affinity.
These three belts are separated from each other
by north-south trending faults. The fault between
the Tokoro and Nemuro belts has a strike-slip
I
I
I
I
I Halaaawa -SW Japan---+l* NE Japan
Hayachine T Z
(3 Oe+ 4 Rni 6 Sn?)
..
10 Sb NE
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302 Y. Isoxaki
\
\ Northeast Japan -\-
ldonappu belt -3 \
\ 14,(Shimanto b.) - , i \
younging polarity --+
Northern Kitakami-
Oshima b.
8. (Mino-Tanba b.) __ East Hokkaido ~
North
10 Sorac hi-Y ezo b.
.(Sanbagawa b.)
~ Yubetsu b. (Shimanto b.) 14.
\'V \
\,Hidaka b?.
\\ (shimanto b.) 100 km
14, 15.
Fig. 11 Geotectonic map of Hokkaido (modified from Kimura unpubl data, 1996) Note the polarity reversal of younging direction in AC units across the Hidaka
Main Thrust which represents a collision suture between the Kurile arc to the northeast Japan arc
nature, and this fault, called the Abashiri tectonic
line, has activated in a right-lateral manner during
the opening event of the Kurile back-arc basin.
Refer to Kimura (1996) for further details.
The main problem in Hokkaido lies in how to
interpret the origin of the parallel-running two
coeval blueschists belts, that is, the Kamuikotan b.
of northeast Japan and the Tokoro belt of East
Hokkaido. The opposite younging polarity in AC
between northeast Japan and East Hokkaido
(Fig. 11) suggests that a secondary collision/
amalgamation has doubled the primarily linear
blueschists belt along the Asian margin. Several
other interpretations were also proposed but a final
agreement has not yet been reached. GEOTECTONIC HISTORY OF THE JAPANESE ISLANDS
The time-space relationships among the orogenic
units described here suggest that the Japanese
Islands have grown oceanward by almost 400 km
across the arc since the mid-Paleozoic. On the basis
of the new geological data and the revised sub-
division of the Japanese Islands, the geotectonic
history of the Japanese Islands is summarized
below. The history began with rifting of a super-
continent in the late Neoproterozoic, and was fol-
lowed by a tectonic inversion shifting from an
extensional (Atlantic-type) regime to a convergent
(Pacific-type) regime around 500 Ma. Since then an
oceanic subduction-related accretion regime has
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Anatomy and evolution of Japanese Islands 303
the core of Japan, which began as a segment of the
Yangtze continental margin. In addition, a frag-
ment of the Sino-Korean continental block also
participated in the orogenic growth of Japan after
the 250 Ma collisional event (Isozaki in press a). persisted, and is responsible for the oceanward
growth of the islands. Figure 12 summarizes this
history from the birth of the islands at ca 700 Ma
to their separation from Asia in the Miocene.
BIRTHPLACE OF PROTO-JAPAN These two continental blocks, particularly the
Yangtze block, have played the most important roles
The Oki belt in southwest Japan and the Southern
Kitakami belt in northeast Japan represents part of in constraining the configuration of the Japanese
Phanerozoic orogen.
I _- -
rifling ,-
I 1.0-07Ga I Asthenospheric mantle
li I1
Break-up of supercontinent Rodinia I and birth of Pacific (750-700 Ma)
(b)
Y~JW~' Passive marain sediments Remnant of orimilive
Onset of initial oceanic subduction
(tectonic inversion) (-500 Ma)
Calc-alkaline 0-eyama fore-arc ophtolite
volcanic arc --, ,, Omanward accretionay growth
in the Paleozoic (400-250 Ma)
lCollision of YangtzeISino-Korea
Hida gneiss
(collision comDlex) Ok Oe Rn Ak 6n.c: Mz V.T Sb SIB
system along continental
Fig. 12 Series of simplified cartoons (a-g) showing 700 million years of geotectonic evolution of the Japanese Islands (a) The birth of Japan along a Proterozoic
rifted continental margin of Yangtze by the break-up of the supercontinent Rodinia and the birth of the Pacific Ocean at 750-700 Ma. (b) After widening of the
proto-Pacific Ocean basin, tectonic inversion occurred to initiate an intraoceanic subduction zone off the rifted Yangtze margin around 500 Ma, leaving a small fragment
of primitiveoceaniccrust (0-eyamaophiolite). (c) Thearc-trench system wasestablished by ca450 Maand gaveacalc-alkalineoverprint on thefore-arcophiolite. (d)
The subduction zone matured to accommodate the oldest AC which corresponds to the protoliths of the 400 Ma high-P/T Renge schists. The high-P/T Renge belt
together with coeval granite belt (fragmented) form the oldest set of paired metamorphic belts. (e) Successive oceanic subduction widened the accretionary edifice
during the Late Paleozoic (the Renge and Akiyoshi belts). On the opposite side of the Yangtze block, another continental block (Sino-Korea) collided and thrust over the
Yangtze at around 250 Ma. (f) Oceanward growth on the Pacific side continued during the Mesozolc and Cenozolc to widen and thicken the accretionary edifice in the
form of subhorizontal piled nappes. The accretionary growth was punctuated by the episodic arrival of mid-oceanic ridge roughly every 100 milion years, which
generated 3 to 4 sets of sandwich structure of high-P/T nappe between low-pressure AC units above and below (e.9. Sangun and Sanbagawa belts; see text for details).
(9) By the back-arc opening of the Japan Sea at 20 Ma, Japan detached from mainland Asia; a continental arc becamean island arc The Hida belt represents a remnant
of the 250 Ma collision suture between the Yangtze and Sino-Korean blocks that partly features an ultrahigh-pressure metamorphic belt.
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304 Y. Isoxuki
The relation of the Yangtze and Sino-Korean
continental margins to protoJapan goes back to
nearly 750-700 Ma when it is believed that the
supercontinent Rodinia started rifting apart (Hoff-
man 1991; Dalziel 1992; Powell et al. 1993). Judg-
ing from the patterns of radial dyke swarms and
rifted basins (Park et al. 1995; Bond et al. 1984),
the breakup of Rodinia was probably triggered by a
superplume rising to the surface from the mantle/
core boundary. Consequently, the proto-Pacific
ocean was born in 750-700 Ma by the rifting of
Rodinia, and several continental fragments, includ-
ing the Sino-Korean and Yangtze blocks were
dispersed in various directions (Fig. 12a). Due to
similarities in Neoproterozoic to Cambrian strati-
graphy, Laurentia (North America) and Australia
are probably the conjugate continental block(s) to
the rifted Yangtze margins (Li et al. 1995). How-
ever, the paleoposition of the Sino-Korean block
with respect to Rodinia remains highly enigmatic.
The birthplace of Japan was probably located some-
where in the northern periphery of Rodinia in mid-
latitudes at about 700 Ma. Most of the dispersed
continental fragments once again assembled to
form another (semi-)supercontinent, Gondwana-
land, about 500 Ma (Hoffman 1991; Dalziel 1992),
however, some large blocks such as Laurentia, Sibe-
ria and Baltica were isolated from the superconti-
nental mass. Likewise, the absence of evidence for
late Neoproterozoic to Cambrian collision in China
suggests that both the Sino-Korean and Yangtze
blocks, including proto-Japan, were also isolated
from Gondwanaland, although Early to mid-Paleo-
zoic faunal provincialism suggests that the Yangtze
block (including proto-Japan) and Australia were
close neighbors (Burrett et al. 1990; Kato 1990).
Due to limited exposure and later tectonic modi-
fication in Japan, there is little evidence to conclude
that continental rifting, such as extensional fault
system, rift-related bimodal volcanism, and rift-
related sedimentary sequences were active at this
time. Stratigraphical and paleontological studies
suggest that the small distribution of Early to
Middle Paleozoic (Ordovician to Devonian) terrige-
nous clastic/carbonate sequences in the peri-
phery of the Hida belt, Hitachi-Takanuki belt,
Southern KitakamiIMatsugataira-Motai belt, and
the Kurosegawa belt all represent remnants of
continental shelf facies accumulated along the
rifted Proterozoic continental margins. The Korean
peninsula and/or mainland China may have pre-
served such features of Rodinian rifting, however,
further structural and stratigraphical analyses are
needed. INVERSION FROM PASSIVE TO ACTIVE MARGIN
With the exception of Precambrian gneissic clasts
in younger sediments, the oldest unit of oceanic
affinity in Japan is the 580 Ma ophiolite in the
O-eyama belt, southwest Japan. Its occurrence in
the periphery of the Oki belt (ancient Yangtze
margin) suggests that this unit is a remnant of
proto-Pacific oceanic crust. The O-eyama ophiolite
and its equivalent in the Miyamori-Hayachine belt
in northeast Japan have a bimodal distribution of
radiometric ages; one at 580 Ma and another at
480-450 Ma (Ozawa 1988; Nishimura & Shibata
1989). Isozaki and Maruyama (1991) explained the
age distribution of the oldest ophiolite in Japan in
the following way.
The tectonic history of the O-eyama ophiolite is
two-fold: (i) following rifting with the emplacement
of nascent oceanic crust, a MORB-like oceanic
crust formed at the proto-Pacific mid-oceanic ridge
around 580 Ma (Fig. 12a); and (ii) these rocks
were then intruded around 450 Ma by calc-alkaline
volcanism of arc affinity when a new intra-oceanic
subduction was initiated (Fig. 12b,c).
O-eyama is regarded as the only example of a
fore-arc ophiolite in Japan, while other ophiolitic
rocks are regarded as accreted fragments of an-
cient seamounts, rises and plateaus (Isozaki et al.
1990b; Kimura & Maruyama in press). Between
580 Ma and 480 Ma, tectonics in proto-Japan
changed dramatically from riftlridge-related ex-
tension to subduction-related compression, and
passive margins changed rapidly to active margins.
This tectonic inversion probably corresponds to
global plate reorganization, in particular to the
opening of the proto-Atlantic (Iapetus) ocean on
the opposite side of the globe. Accretion of the
O-eyama ophiolite can be explained by either of the
following two mechanisms: (i) initiation of a land-
ward dipping subduction zone within the primary
oceanic crust (see Fig. 12b,c); or (ii) collision of an
island arc system from the ocean side and a
reversal of subduction polarity.
ACCRETIONARY GROWTH OF THE JAPANESE ISLANDS
After the tectonic inversion around 500 Ma, proto-
Japan began a state of accretionary growth that
persists today. Within 50 million years after initi-
ation of intra-oceanic subduction, the arc-trench
system featured AC, high-PIT schists (metamor-
phosed AC), and granitic batholiths (Fig. 12d,e)
that is, in the Renge belt plus Kurosegawa belt
(=tectonic outlier of the former) in southwest
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Amtomy and evolution of Japanese Islands 305
Sino-Korean and Yangtze blocks collided along the
Qinling-Dabie suture, generating an ultrahigh-
pressure metamorphic belt, best preserved in the
Dabie mountains and Shandon province in China
(Wang et al. 1989; Maruyama et al. 1994; Cong &
Wang 1995). The region of ultrahigh-pressure
metamorphism appears to be restricted to the
central part of the 3000 km long suture, suggest-
ing that a promontory collision may have locally
generated ultrahigh-pressure conditions (Isozaki in
press a). East of the Shandon peninsula, the coeval
250 Ma regional metamorphism is detected solely
in the Hida and Hitachi-Takanuki belts in Japan;
they are characterized by medium-pressure-type
metamorphism. Although these belts lack ultra-
high-pressure metamorphism, from a geographical
viewpoint (Fig. 13), they may also correspond to an
eastern extension of a collision-related metamor-
phic belt (Sohma et al. 1990), including the Imjin-
gang or Ogchon zones in Korea (Ernst et al. 1988;
Lan et al. 1995). The Hida belt and its equivalent
occur at the eastern terminal of the suture. Data
from this belt indicate a decrease in grade and
magnitude of a collision-related metamorphic belt.
After 250 Ma, accretionary growth in Japan radi-
ated from a core composed of the amalgamated
Sino-Korea plus Yangtze blocks. Thus while minor
in extent, the Japanese Islands do contain elements
of a non-accretionary, continent-continent collision-
type orogen. Japan, and in Southern Kitakami and Matsuga-
taira-Motai belts in northeast Japan. In particular,
the oldest AC in Japan is the protolith of the 450-
400 Ma high-P/T schists. The high-P/T schists and
coeval granitic rocks are elements of a paired
metamorphic belt (with a high-P/T belt on the
ocean side and a low-P/T belt on the continent
side; Miyashiro 1961).
Following the oldest 450 Ma unit, Late Paleo-
zoic, Mesozoic and Cenozoic AC were formed
through subsequent subduction. At least several
major oceanic plates have subducted beneath the
Yangtze margin, leaving more than 10 distinct AC
belts. Numerous oceanic fragments derived from
subducted oceanic plates, including deep-sea sedi-
ments and seamount-derived basaltslreef lime-
stone, were accreted to Japan. Details of the
accretion processes during the Permian and Juras-
sic periods are reported in Isozaki (in press a,b).
Accretionary growth apparently has been not
continuous. Including the youngest AC now under
construction along the Nankai trough off southwest
Japan, total accretionary growth is nealy 400 km
in across-arc width (Fig. 12f), not taking into
account the material loss by subduction-erosion.
Thus the overall AC-dominated orogen in Japan
has grown oceanward for almost 400 km during
the 450 million years (-100 km per 100 million
years).
As all of the AC units in Japan were formed in
situ by subduction along the Yangtze (South
China) continental margin, they are autochthonous
to Asia, with the exception of small oceanic frag-
ments peeled off from subducting oceanic crust and
accreted landward into AC (Isozaki et al. 1990b).
It thus appeared that the Japanese islands do not
represent a collage of ‘suspect or exotic terranes’
that had existed prior to subduction and accretion
processes (Coney et al. 1980).
It was a mere 20 million years ago when the
Japanese Islands obtained their present configura-
tion through back-arc spreading (Fig. 12g). How-
ever, the accretionary growth of the Japanese
Islands will likely continue until other continental
blocks, such as Australia or North America, collide
against Asia to form a future supercontinent
(Maruyama 1994).
REMNANT OF CONTINENT-CONTINENT COLLISION
While most of the geotectonic units in Japan are
the results of oceanic subduction, the Hida belt and
the Hitachi-Takanuki belt are remnants of
continent-continent collision. At about 250 Ma, the DISCUSSION
The newly revised geotectonic subdivision and re-
constructed tectonic history of the Japanese Is-
lands clarifies the properties of the Cordilleran-type
orogenic growth. Five tectonic features are de-
scribed following these observations in Japan. They
are: (i) step-wise growth of AC; (ii) subhorizontal
nappe structure; (iii) downward younging polarity;
(iv) a tectonic sandwich of high-P/T metamorphic
units; and (v) secondary tectonic modifications.
These may represent the principal characteristics
of subduction-related orogenic belts in general.
STEP-WISE GROWTH OF AC UNITS
The 450 million-year-old subduction-related history
of the Japanese Islands is characterized by the
intermittent formation of AC, clearly detected by
OPS analysis (Fig. 7). Several intervals lack AC
(Carboniferous-Early Permian, Early-Middle Tri-
assic, and late Early Cretaceous), and this empha-
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306 Y. Isoxaki
Jurassic AC
sizes the overall zonal (piled nappe) arrangement
of AC in Japan (Figs 5,6,8-11). Secular changes in
relative plate motion between the Asian continent
(or its precursory fragment) and subducted oceanic
plates may be responsible for such episodicity
because AC cannot form from a highly oblique
subduction or in an along-arc strike-slip regime
(Maruyama & Sen0 1986).
On the other hand, extensive modern AC can be
constructed where a large amount of sediments are
supplied and the subduction rate is moderately low.
No accretion or tectonic erosion (subduction-
erosion) occurs when a trench is starved of sedi-
ments or the subduction rate is too high (von
Huene & Lallemand 1990; von Huene & Scholle Fig. 13 Tectonic framework of 250 Ma
East Asia showing a continent-continent col-
lision orogen between the two Precambrian
cratons that is the Sino-Korean and
Yangtze blocks (modified from lsozaki &
Maruyama 1991) The collision suture fea-
turing the 250 Ma ultrahigh-pressure rneta-
morphic belt of the Qinling-Dabie zone in
China extends eastward to the Hida belt and
Hitachi-Takanuki belt in Japan character
ized by the 250 Ma medium-pressure meta-
morphism The pathway of the suture in the
Korean peninsula is still controversial as two
alternative interpretations are proposed, one
along the Ogchon zone and the other along
the Imjingan zone
1991). Thus the apparent episodicity in the forma-
tion of AC in Japan may also indicate the episodic
preservation of AC rather than formation. None-
theless such episodic growth of AC is a primary
characteristic of oceanic subduction-related oro-
gens that persist for 100 million years or longer.
The formation of high-PIT metamorphic and
coeval granitic belts also appears episodic as these
occur in highly restricted time intervals. As they
occur roughly every 100 million years, their forma-
tion may also be episodic and/or periodic.
In contrast, the classic concept of the Cordil-
leran-type orogeny proposed by Dewey and Bird
(1970) assumed a steady-state orogenic process,
including the simultaneous formation of subduction
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Anatomy and evolution of Japanese Islands 307
Subhorizontal piled nappe structures generally
characterize continent-continent collision-type oro-
gens, like the European Alps or the foreland fold-
and-thrust belt in the Appalachian and Canadian
Rocky mountains. Development of similar struc-
tures in an arc-trench setting dominated by sub-
duction tectonics is noteworthy and should be
tested in other AC-dominated orogens older than
100 million years. complexes ( = AC), granitic batholiths, regional
metamorphic belts, and relevant geologic struc-
tures. Later publications, h.owever, emphasized the
non-steady-state nature of the arc-subduction zone
setting (Dewey 1980). Although various interpre-
tations are given, causes for the episodicity in
subduction-related orogeny has not been well clar-
ified. A possible cause for the episodic exhumation
of high-PIT metamorphosed AC and coeval gra-
nitic activity associated with low-P/T regional
metamorphism will be discussed later.
SUBHORIZONTAL PILED NAPPE STRUCTURE OF AC
Most of the geotectonic boundaries between adja-
cent AC units in southwest Japan are low-angle
faults, and therefore the AC units are considered to
be piled nappes (Fig. 6). Recent microfossil/
chronometric mapping documented the occurrences
of regional klippes and windows of these nappes
even in thickly vegetated areas. These observations
indicate that the subhorizontal piled nappe struc-
ture governs the fundamental tectonic framework
of the 450 million-year-old orogen in southwest
Japan (this study: see also Hara et al. 1977;
Charvet et al. 1985; Faure 1985). Primary oro-
genic structures in the Ryukyus and northeast
Japan have not yet been fully mapped but a
subhorizontal piled nappe structure controlled by
subsurface blind thrusts has been predicted
(Tazawa 1988; Isozaki & Nishimura 1989;
Fig. 10).
To a first approximation, such a subhorizontal
piled nappe structure is consistent with subduction-
related deformations, particularly to the subhori-
zontal shortening caused by underplating through
which new materials are added to the sole of the
previously formed accretionary wedge through
step-wise activation of a subhorizontal decolle-
ment. The size of an individual ancient AC nappe is
nearly 200 km in width across the arc, similar to
the size of the widest modern AC wedge (von
Huene & Scholle 1991).
The documentation of a predominant subhorizon-
tal structure in Japan is contrary to the traditional
view that vertical tectonics dominated horizontal
tectonics. There are, in fact, some along-arc verti-
cal faults of strike-slip nature (Neo-M.T.L.,
T.T.L.), but most of them became active in the
Cenozoic and were driven by microplate activities.
The total amount of displacement along the verti-
cal fault, however, is too small to account for the
present zonal arrangement of belts in Japan that
extend along the arc for more than 1000 km. DOWNWARD YOUNGING POLARITY IN AC NAPPES
The along-arc zonal arrangement of these ancient
AC is most clearly demonstrated in southwest
Japan (Fig. 14a). Remarkably, these AC show an
oceanward younging polarity in map view without
windows and klippes. In addition, ages of regional
metamorphic belts and granitic batholiths also sug-
gest an oceanward younging polarity. This appar-
ent polarity is a function of the piled nappe struc-
ture of the AC (Figs 6,12f). Downward younging
polarity is also recognized within an individual AC
nappe; this was clearly demonstrated in the Juras-
sic AC, which comprises six or more subnappes
(Isozaki in press b). Such oceanward and down-
ward younging polarity is consistent with the
growth patterns of modern AC.
Although detailed information is scarce, the
Ryukyus and northeast Japan plus East Hokkaido
also appear to display younging polarity. The
present zonal arrangement in northeast Japan,
however, does not fit with the oceanward younging
polarity in southwest Japan (Fig. 14b). This is due
to secondary strike-slip faulting related to the
opening of the Japan Sea in the Miocene which
modified the primary orogenic configuration.
TECTONIC SANDWICH OF HIGH-P/T AC UNIT
There are three geotectonic units composed of
high-P/T schist in the Japanese Islands formed at
around 450-300 Ma, 200 Ma, and 100 Ma. The
450-300 schists can be further subdivided into two
distinct units. All of these high-P/T units also
occur as subhorizontal nappes (Fig. 6). The most
striking feature of the high-P/T nappes is their
sandwich-like structure, where a high-PIT nappe is
tectonically interleaved between two unmetamor-
phosed AC nappes (Maruyama 1990; Isozaki &
Maruyama 1991). The nappes composed of high-
P/T metamorphosed units can be traced laterally
for more than 500 km along the arc, even though
they are usually thinner than 2 km. For example,
the 100-70 Ma Sanbagawa schists (Sb) occur as a
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308 Y. Isoxaki
SW-Japan / Ryukyus NE Japan
Sino - Korean crafon
Precambrian craton collision zone
ophiolite / high-P/T metamorphosed AC non- to weakly metamorphosed AC Shn
\
\
low-PrT metamorphosed part
Fig. 14 Cartoon showing stepwise oceanward growth of southwest Japan with a high-P/T sandwich structure after removing windows and klippes (see text) in
a map view and a contrasting feature in northeast Japan which was secondarily modified by strike-slip shredding lo obscure the primary structure
nappe sandwiched between the overlying Jurassic
AC (MT) and the underlying Cretaceous AC (Sh)
(Kawato et al. 1992; Sasaki & Isozaki 1992).
There are high pressure gaps between the high-
P/T nappe in the middle of the sandwich and the
adjacent AC nappes. These gaps, which are greater
than several kilobars in pressure, correspond to
10-20 km of crustal thickness. In order to preserve
these pressure gaps without metamorphic anneal-
ing, the tectonic insertion of the high-P/T nappe
must have been rapid and caused by tectonic
exhumation of a high-PIT nappe into a low-
pressure domain. In fact, all of the high-P/T
nappes are bounded by a set of subhorizontal faults
along their top and bottom surfaces, however, the
sense of dislocation between these faults is oppo-
site. To compensate for the pressure gap between
the hanging wall and foot wall, the upper fault is
normal while the bottom fault is reverse. Synchro-
nous activation of such paired faults is believed to
result in the insertion of the thin high-P/T nappe
into the low-pressure domain.
These observations strongly suggest that the
exhumation of the high-PIT nappes was tectonic and episodic rather than caused by steady-state
processes wherein buoyant uplift of a metamorphic
domain together with over- and underlying un-
metamorphosed units was proposed. This episodic
tectonic exhumation mechanism called ‘wedge ex-
trusion’ was first suggested by Maruyama (1990;
see also Maruyama in press). A very similar model
was also proposed for Himalayan medium-pressure
gneisses by Burchfiel et al. (1992). The downward
younging polarity among AC is not disturbed by
the intermittent intercalation of high-P/T nappes.
This suggests that subduction-driven burial of the
protolith AC, high-PIT metamorphism, and tec-
tonic exhumation of metamorphosed AC all oc-
curred in the same structural horizon, likely along
the Wadati-Benioff plane. The large-scale sandwich
(containing a thin well-done steak) with downward
younging polarity cannot be formed through any
other known exhumation process (Suppe 1974;
Cloos 1982; Platt 1986).
Based on radiometric ages determined by various
methods, the timing of metamorphic peak temper-
ature and the subsequent cooling history of high-
P/T units in Japan (Itaya & Takasugi 1989; Nishi-
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Anatomy and eTolution qf Japanese Islands 309
gests an eastward along-arc younging polarity for
these intrusions (Nakajima et al. 1990; Kinoshita
1995). This observation is consistent with the
oblique subduction of the Kula/Pacific ridge be-
neath Asia as mentioned above (i.e. northeastward
passage of the TTR triple junction of the Eurasia/
Kula/Pacific plates, because the subducted ridge-
related slab window may have been a heat source).
The coincident timing of high-P/T nappe exhuma-
tion and granite belt formation suggests a causal
relationship to ridge subduction. Further research is
needed to elucidate the mechanism of the subhori-
zontal ejection of high-PIT nappes under a buoyant
subduction regime induced by the movement of
young oceanic crust. The preservation of high-PIT
units in the fore-arc is possible even under high heat
flow from the subducted ridge if the exhumation of
the high-PIT nappe preceded the arrival of the
ridge-crest at the trench. The metamorphic/cooling
ages of the high-P/T unit (100-70 Ma) are slightly
older than the subduction timing of the brand new
oceanic plate (ca 75 Ma), and this may support the
interpretation provided above.
Prior to the 100-70 Ma event, exhumation of the
high-PIT nappe occurred probably three times in
the 600 million year history of the Japanese Islands
(Fig. 15). Such episodic occurrences of the high-
PIT sandwich structure suggests that these were
also consequences of episodic ridge-subduction in
the Paleozoic and early Mesozoic (Fig. 16). The
distribution of pre-Cretaceous granites in Japan is
highly limited; most of them are fragmented.
Nonetheless the ages of the granites that are
preserved suggest episodic formation of a granitic
belt in the intervals of 450-400 Ma, 350 Ma, and
250 Ma.
If those three pre-Cretaceous geologic episodes
are regarded as the result of episodic ridge subduc-
tion, the tectonic history of the Japanese Islands mura et al. 1989; Shibata & Nishimura 1989;
Takasu & Dallmeyer 1990), the exhumation of
high-P/T nappes occurred episodically. It is note-
worthy that these episodic events apparently coin-
cide with the formation of granitic batholith in
low-P/T metamorphic belts on the continent side
(Isozaki & Maruyama 1991; Fig. 15). The best
example of such paired metamorphic belts, de-
scribed by Miyashiro (1961), is the 100-70 Ma
high-P/T Sanbagawa belt on the trench side and
the low-P/T Ryoke belt on the arc side of south-
west Japan.
The formation of the Cretaceous paired belts was
probably related to subduction of the Kula/Pacific
ridge along the Asian margin. An observation on
contemporary OPS in mid-Cretaceous AC in the
Shimanto belt in southwest Japan (Taira et al.
1988) suggested a progressive decrease in the age
of the subducted slab in the Late Cretaceous
(Fig. 16). At about 75 Ma, an AC was formed by
subduction of a brand new oceanic plate (a mid-
oceanic ridge per se).
In addition, based on hot spot tracks, the recon-
structed paleo-plate motion in the Pacific domain
over the last 150 million ,years indicates that the
Kula-Pacific mid-oceanic ridge subducted obliquely
beneath the eastern Asian continental margin
around 100-70 Ma (Engebretson et al. 1985).
Thus a TTR (trench-trench-ridge) triple junction
may have migrated from southwest to northeast
off-shore of Japan.
The phenomenon of ridge-subduction and its
tectonic effects on the continental margin have long
been discussed (Pitman & Hays 1968; Wilson
1973; Uyeda & Miyashiro 1974); however, its
tectono-thermal effects and orogenic consequences
are still controversial (Farrar & Dixon 1993; Thor-
kelson 1994). A recent compilation on the age of
Mesozoic granite emplacement in East Asia sug-
Stepwi$e oceanward propaRation of BS belt and granite belt Fig. 15 Simplified diagram of geotectonic
profile of the Japanese Islands (modified
from lsozaki & Maruyama 1991) Note the
oceanward and tectonically downward N
younging growth of AC nappes incluijing
high-P/T blueschist (BS) nappe Formation
of the sandwich structure of BS nappe occurs
intermittently roughly every 100 million
years, accompanying coeval granite batho-
lith belt with low-P/T metamorphic belt
which is usually positioned -100-200 km
continentward from coeval BS belt The iriter-
mittent oceanward propagation of the BS belt
and granite belt appears to have occurred
when mid-oceanic ridges episodically col- b
High-PIT metamorphic AC (BS) nappe
Low-pressure AC iiappe
Granite intrusron /
tided and subducted beneath proto-Japan
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
310 Y. Isoxaki
Age (Ma)
Formation of
high-PIT sandwich
(tectonic exhumation)
Age of subducted 400 300 200 100 0 ,
I
450-40018 330-280Ma 230-210~a 100-70Ya
Renge I1 Sang"" Sanbsgawa 'engel * * * *
-~ I
I/ --0 1
,' (\ I, io
15 slabdeduced from
OPS of AC
interacted oceanic plate
Ridge subduction
can be simplified to 450 million years of oceanward
accretionary growth punctuated four times by
ridge subduction. For each event, granites and
associated low-P/T metamorphic belts formed on
the continent side of Japan, probably separated by
200 km or more from the coeval high-PIT schist
belt (Fig. 15). The nearly 100 million year period-
icity in sandwich formation may correspond to the
consumption of a major oceanic plate dissected by
mid-oceanic ridges on both sides. The Japanese
Islands experienced subduction of at least five
major oceanic plates (from young to old, Pacific,
Kula-Izanagi, Farallon, an unnamed older oceanic
plate, and the plate which originated during initial
rifting; Fig. 16).
The anatomy of the subduction-related orogen in
Japan and the tectonic interpretation of these
structures strongly contradict earlier views of
'Cordilleran-type' orogeny. The most significant
difference lies in the recognition of episodicity in
orogenic activity related to ridge subduction. Pre-
vious models for exhumation of high-P/T units
assumed a steady-state and long-term process re-
lated to subduction of the normal ocean floor. For
such an oceanic subduction-related orogenic pro-
cess involving episodic culmination by ridge sub-
duction, Isozaki and Maruyama (1 99 1) introduced
a new term, 'Miyashiro-type orogeny'. The name is
given after Akiho Miyashiro's outstanding contri-
butions to subduction zone tectonics, particularly
the first perception of paired metamorphic belts
(Miyashiro 1961) prior to the birth of plate tecton-
ics in the mid-l960s, and a keen perspective on
ridge subduction and relevant geologic phenomena
in the 1970s (Uyeda & Miyashiro 1974). ,' '\\ 7
- LOO ,' ?
~ 150
-- 200Ma-
we7 Unnamed Farallon $1 Izanagi ,,*" Kula $1 Pacificp
?nn 77n
SECONDARY MODIFICATION: MICROPLATE TECTONICS
The primary orogenic structures of the Japanese
Islands are controlled by the spatial arrangement Fig. 16 Timetable of ridge subduction and
orogenic culmination in Japan indicated by
formation of the high-P/T sandwich structure,
age of subducted oceanic slab deduced from
OPS analysis and paleoplate interaction
based on hot spot track analysis (modified
from Isozaki & Maruyama 1991) The names
for oceanic plates that interacted with Japan
are adopted from the reconstructed plate in-
teraction in the Pacific domain by Engebret-
son eta/ (1985) Note the temporary coinci-
dence among the zero age of subducted
oceanic plate at 75 Ma exhumation of the
high-P/T Sanbagawa metamorphic belts, and
passage of KulaiPacific ridge off Japan
of main components (i.e. ancient AC, regional
metamorphic rocks and granitic batholiths).
Subduction-related orogens in Japan are character-
ized by subhorizontal piled nappes of AC with
downward younging polarity, and by the episodic
occurrence of a sandwich structure of high-PIT
nappes and unmetamorphosed AC. The primary
orogenic features, however, were modified or de-
stroyed in some cases by secondary tectonic pro-
cesses. In particular, microplate-related tectonics
has the profound ability to secondarily reorganize
primary structures (Miyashiro 1982). In Japan,
there are three important secondary tectonic pro-
cesses (Fig. 17; fore-arc sliver movement, back-arc
basin opening, and arc-arc collision) that are all
orogenic manifestations of microplate activities
(Isozaki 1989; Isozaki & Maruyama 1991).
A fore-arc sliver is a decoupled microplate of a
frontal arc driven by a strike-slip component of the
oblique subduction of an oceanic plate. In the
Japanese Islands, there are three active fore-arc
slivers (the South Ryukyu sliver, Nankai sliver in
southwest Japan, and Kurile sliver in East Hok-
kaido; Fig. 1). Along-arc movement of fore-arc
slivers can create three distinct features that de-
stroy the primary orogenic edifice (i.e. along-arc
strike-slip fault on are-side margin, and across-arc
compressional structure in front, plus an exten-
sional one in rear of the sliver). The best example
of an along-arc strike-slip fault is the Quaternary
Neo-M.T.L. in the Kii peninsula and on Shikoku
Island that demonstrates a remarkably linear sur-
face trajectory for more than 500 km (Figs 3,5).
This right-lateral strike-slip fault was driven by
the westward movement of the Nankai fore-arc
sliver (Fig. 17b), cutting the older low-angle fea-
ture of the Tertiary paleo-M.T.L. (Isozaki 1989;
Yamakita et al. 1995; Fig. 6). This westward sliver
translation is also associated with across-arc com-
pression along the Bungo Strait between Shikoku
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Anatomy and evolution of Japanese Islands 311
(a) Arc-arc collisian (b) Fore-arc sliver
Philippine Sea Plate
(c) Opening of back-arc basin I’ A\
Fig. 17 Three representative modes of secondary modification of primary orogenic structure related to microplate tectonics (modified from lsozaki & Maruyama
1991). (a) Formation of syntaxis in the older accretionary orogenic system and accretion of exotic arc crust occur at an arc-arc collision front around the Izu
peninsula where the Izu arc penetrates almost perpendicularly to the Southwest Japan arc (b) Formation of along-arc strike-slip fault (Neo-M T.L ) and across-arc
extensional and compressional structure by along-arc movement of the Nanakai fore-arc sliver (c) Formation of along-arc extensional structure (Miocene rifted
basins with bimodal volcanism), compressional structures (thrusting along paleo-M.T L ) and the strike-slip fault (T T.L.) occurred when the Japan segment
detached from mainland Asia through the rifting-opening of the Japan Sea. Note these Neogene to Quaternary structures have modified considerably the pre-existing
major structures of the ca 450 Ma accretionary orogen
and Kyushu Islands, and across-arc extension in
Ise Bay. Across-arc compression related to a fore-
arc sliver is best observed in the elevated Hidaka
mountains in central HoIkkaido. Kimura (1985)
explained the exposure of lower crustal rocks of
the Hidaka belt (Komatsu et al. 1989) as a tectonic
manifestation of westward frontal collision of the
Kurile fore-arc sliver to northeast Japan. Seismic
reflection research (Ikawa et al. 1995) recently
documented the crust-cutting detachment surface
that dips eastward from the Hidaka main thrust
(Fig. 11).
Back-arc spreading is another tectonic process
that has occurred frequently in the western Pacific
since the latest Mesozoic, and is also a powerful
modifier of primary orogenic structures in Japan.
For example, when the Japan Sea opened in the
Miocene, it split pre-existing orogenic structures
into several blocks (Figs 9,10,14). Although a cer- tain amount of rotation was involved, the opening
of the Japan Sea basin has been attributed to
dislocation of a pair of north-south running strike-
slip faults along the basin margin (Fig. 17c). The
chaotic alignment of belts in northeast Japan is
primarily due to the series of left-lateral strike-
slips along the eastern margin of the Japan Sea.
The T.T.L. presently dividing southwest Japan and
northeast Japan, and the parallel Futaba and Hata-
gawa faults are typical examples of an eastern
margin strike-slip fault that obliquely dissects an
older along-arc zonal arrangement. In contrast, a
segment of the right-lateral strike-slip fault mark-
ing the western margin in east Korea was recently
noted (Yoon & Chough 1995). In addition, a do-
main of compression in the fore-arc may have been
associated with back-arc spreading (Fig. 17c). It is
difficult to explain the side-by-side juxtaposition of
the low-P/T Ryoke metamorphic belt and high-PIT
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312 Y. Isoxaki
Sanbagawa metamorphic belt in modern southwest
Japan (Fig. 5). These belts should have been sepa-
rated from each other by at least 100-200 km
when they formed in the Late Cretaceous arc-
trench system (Fig. 18a). The origin of the low-
angle Paleo-M.T.L. between these two belts
(Fig. 6) may be related to development of a hori-
zontal detachment surface in the arc crust and to
trenchward dislocation of the upper crust along this surface, probably by back-arc spreading
(Isozaki & Maruyama 1991; Fig. 18b). For further
details on the opening of the Japan Sea, refer to
Jolivert et al. (1994), Otofuji (1996) and Yamash-
ita et al. (1996). The present-day opening of the
Okinawa trough (Kimura et al. 1988) is still in a
nascent stage of back-arc spreading (Fig. 1). This
extensional regime appears to propagate north-
ward into mid-Kyushu (note the Beppu-Shimabara
6 c usually 100-200 km (100-701 a? ca 50 km ___b.
Akryorhr-Sangun A B C
pre-Juraralc
S posslble
external llmlt of
volc. tront trench
- -
high-P/T Sanbagawa met.
Mlno-lanba
Juraralc complex
100 km
subducted
Kula-Paclflc
rldge
Orlglnal conflguratlon of the Cretaceous
palred metamorphlc belta in SW Japan p i~
- n'B' = Intra-arc shortenlng 1 = dlslocatlon 01 delslchmenl Iault I 130-40 Ma] fore-arc contraction 1
A', B' C'
100 Ma trench 100 Ma voic front
trench Paleo-MTL volcanlc [rant
Shlmanto
Palaogana complex back -arc
extension
mlsalnp volume of the Ryokc fore-arc rcplon
= accreted volume 01 the Shlmanto complex
Secondary juxtapositlon of
hlgh-P/T and low-P/T motamorphlc belts
Fig. 18 Schematic cross-section showing the primary configuration of the Cretaceous arc-trench system in southwest Japan (a) and secondary modification
involving intra-arc shortening associated with back-arc (Japan Sea) spreading (b) Note the juxtaposition of the Cretaceous paired metamorphic belts induced by
the trenchward translation of upper arc-crust along an intracrustal horizontal detachment surface
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Anatomy and evolution of Japanese Islands 313
tica, and eastern Australia, as the Paleo-Pacific
ocean floor was subducted along these continental
margins after the 750-700 Ma breakup of Rodinia.
Reconnaissance studies (Sedlock & Isozaki 1990;
Isozaki et al. 1992; Isozaki & Blake 1994) includ-
ing a compilation of previous works for the
Klamath-Franciscan belt in California (Isozaki &
Maruyama 1992; Maruyama et al. 1992) suggests
fundamental similarities between the geotectonic
evolution of California and that of Japan. It ap-
pears promising to extrapolate the above-described
new aspects and analyzing schemes for ancient AC
to other orogenic belts of different time-space
co-ordinates in the Earth’s history, including the
Precambrian. graben with active Aso and Unzen volcanoes),
accordingly, Kyushu Island, now composed of sev-
eral geotectonic units, may eventually split into
two.
Arc-arc collisions are the third process that can
modify the primary geotectonic structures in Ja-
pan. The ongoing collision of the Izu-Bonin arc
against the southwest Japan arc is a typical exam-
ple of this process (Fig. 17a). The northward buoy-
ant subduction of an intra-oceanic arc can indent
into a pre-existing orogenic structure, leaving a
clear V-shaped mark called ‘orogenic syntaxis’, in
the west of Tokyo (Fig. 4). In addition, the accre-
tion of arc crust basement has been achieved in a
step-wise manner, adding a significant volume in
central Japan (Taira et al. 1989). Several tectonic
blocks of intra-oceanic arc origin around Mt. Fuji
(Amano 1986) thus represent bona fide ‘allochtho-
nous or exotic terranes’ in Japan.
CONCLUSION
The Japanese Islands represent a segment of a
Phanerozoic subduction-related orogen developed
along the western margin of the Pacific Ocean.
Significantly, this orogen has grown oceanward
nearly 400 km during the last 450 million years.
The anatomy of the Japanese orogen, best pre-
served in southwest Japan, is characterized by a
subhorizontal piled nappe structure, which involves
multiple AC nappes including those of high-P/T
metamorphosed AC. Orogenic growth is unusual in
that it includes: (i) step-wise growth of AC units;
(ii) tectonically downward younging polarity; and
(iii) intermittent sandwich structure of high-PIT
nappes. Episodic subduction of mid-oceanic ridges
( =migration of TTR triple junctions) appears to
explain the episodic exhumation of high-P/T
nappes and the formation of granite/low-P/T
metamorphic belts every 100 million years. Micro-
plate tectonic processes such as the movement of
fore-arc slivers, back-arc spreading and arc-arc
collisions, secondarily modified the primary struc-
ture of Japan.
In order to establish a general model for
subduction-related orogenic processes, the new
geotectonic model for the origin of Japan needs to
be tested on other orogens formed in similar tec-
tonic environments. For example, the occurrence of
a subduction-related orogen characterized by a
similar anatomy would be expected in Circum-
Pacific Phanerozoic orogens in western North
America, western South America, Western Antarc- ACKNOWLEDGEMENTS
The author would like to thank S. Maruyama, Y.
Nishimura, T. Itaya, T. Matsuda and many stu-
dents from Yamaguchi University for a decade-
long discussion in various tectonic aspects of the
Japanese Islands. Sincere thanks are also due to
A. J. Kaufman, S. Maruyama, G. Kimura, and
L. Ivany who constructively reviewed the manu-
script. P. F. Hoffman gave valuable comments on
the neo Proterozoic distribution of continental
blocks.
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APPENDIX: HISTORICAL REVIEW ON STUDIES OF
OROGENY AND GEOTECTONIC SUBDIVISION OF THE
JAPANESE ISLANDS
In order to establish a geotectonic subdivision for a
particular region, conventional field mapping on a
regional scale is essential, and this usually requires
painstaking work by numerous geologists over an
extended period of time, particularly in thickly
vegetated regions like the Japanese Islands that
are at a humid mid-latitude. In the 100-year his-
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have been made to synthesize all available geolog-
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Terrane. In Ichikawa K., Mizutani S., Hara I., Hada S.
& Yao A. eds. Pre-Cretaceous Terranes of Japan,
pp. 185-201, Publication of IGCP #224, Osaka.
Japanese Islands can be roughly divided into four
stages, based on the orogenic concept that gov-
erned the understanding of contemporary geolo-
gists; these are: (i) the pre-geosyncline concept
stage (1865-1940); (ii) the stage of importing the
geosyncline concept (1941-1955); (5) the stage
of popularization of geosyncline-based orogeny
(1956-1975); and (iv) the stage of plate tectonics-
based orogeny (1976-present).
The first stage is represented by activities of E.
Naumann, B. Lyman and other foreign geologists
who imported modern geology as well as other
sciences to Japan immediately after the Meiji revo-
lution in 1868, after two century-long diplomatic
isolation. During this stage, foreign and domestic
geologists recognized fundamental structures of
the Islands, particularly major tectonic boundaries
such as the Median Tectonic Line (M.T.L.) and
Itoigawa-Shizuoka Tectonic Line (I.-S. T.L.); how-
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Anatomy and evolution of Japanese Islands 319
the importance of distinguishing orogenic phase in
H. Stille’s sense. His summary was the first appli-
cation of the then world-popular geosyncline con-
cept to a terra incognita named Japan. His
achievement, however, marks a prime milestone in
geological studies in Japan and his subdivision
provided a foundation for later works. One point to
note is his emphasis on nappe structures through-
out southwest Japan, as this idea was revived in
the 1990s in the plate tectonic framework.
The third stage was governed by geologists of
the post-World War I1 generation who strongly
criticized Kobayashi’s summary. They emphasized
the higher resolution and accuracy of their data set
through detailed mapping and the advantage of
their new tectonic models (Yamashita 1957: Minato
et al. 1965: Ichikawa et al. 1970), however, their
perspective was more or less the same as that of
Kobayashi, that is, popularization of the classic
geosynclinal concept with minor modification in the
context of regional geology. Although the world
according to fixism was still there, the quality of
the regional geological description of the Japanese
Islands improved much during this stage, with the
use of megafossil dating, particularly fusulinid
biostratigraphy. Regional distribution of geotec-
tonic units and locations of mutual boundaries were
clarified, and the geotectonic subdivision made in
this stage appears very similar to the present one
in map view except for ‘vertical ’ major boundary
faults (Fig. 19b). Accordingly, sporadically exposed
older mid-Paleozoic rocks, including high-grade
metamorphic rocks and serpentinites, were all ex-
plained as geosynclinal basement rocks squeezed
out from deeper levels along putative ‘crust-
penetrating vertical faults’. In the early 1970s,
several avant-garde geologists from Japan were
the first to propose plate tectonic interpretations
for the evolution of the Japanese Islands, with
special emphasis on subduction-related tectonics
(Matsuda & Uyeda 1971: Uyeda & Miyashiro
1974). The majority of scientists, however, still
held conservative understandings linked to the geo-
syncline concept, and the geotectonic subdivision
was not revised from the plate tectonic viewpoint.
The most significant reform in the geotectonic
subdivision occurred in the fourth stage. Nearly a
decade after the fundamental construction of plate
tectonics in the late 1960s, younger Japanese
geologists started to prefer mobilism rather than
fixism. Down a long and winding road with many
arguments, including the conversion from a geo-
syncline world to a plate tectonics world, both in
individuals and in society (Kanmera 1976), re- ever, the overall geotectonic subdivision of the
Islands was still a rough sketch. In hindsight, this
era up to the 1930s in Japan may be called a time
of fundamental ‘find-and-describe’ in preparation
for the following stage of importing the concept of
the geosyncline.
During the second stage, T. Kobayashi was the
first Japanese geologist to synthesize a grand view
of the geotectonic evolutim of the Islands on the
basis of the stratigraphic and megafossil age data
(Kobayashi 1941). He recognized most of the im-
portant geotectonic units currently known, but his
subdivision (Fig. 19a) was strongly biased by the
geosyncline concept. In particular, he emphasized
Akiyoshi Orogen Sakawa Orogen
~~~~~~
MTL
IctoScalel Opa Nappa
:b) lchikawa era/ 1970
N Shimanto Orogen S ~- Honrhu Orogcn ___
Cp Mz MT Ry Sb Ch Hdk Mr
km MTL
[i 30 . ,50 km, ,
Fig. 19 Classic examples of geotectonic map and profile of the Japanese
Islands (modified from (a): Kobayashi 1941, (b): lchikawa et a/. 1970).
Compare these classic geotectonic subdivisions and profiles backgrounded by
geosynclinal viewpoint with the current version shown in Figs 3,4 Before
1975, major geotectonic units in Japaii were all explained as Precambrian
sialic basement and overlying Paleozoic to Mesozoic geosynclinal sediments.
Hd. Hida b , Ok: Oki b., Cg: Chugoku b., Mz, Maizuru b., MT: Mino-Tanba b.;
Ry: Ryoke b.; Sb: Sanbagawa b.; Ch: Chichibu b.; Hdk. Hidakagawa b. ( = Sh:
Shimanto b.); Mr. Muro b.( = NK. Nakamura b.); Km: Kitakami marginal b.;
KI: Kitakami b ~ Sm. Soma b.; Ab: Abukuma b., As: Ashio b.; Jo: Joetsu b.
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
320 Y. Isoxaki
gional geologic information was constantly accu-
mulated on a nationwide basis. The great reform
came in two waves, the first one in the early 1980s
and the second in the early 1990s.
The first wave was brought by two factors com-
bined: (i) the remarkable enhancement in micro-
fossil (conodont and radiolaria) biostratigraphy, in
particular the high resolution dating of the pre-
Cretaceous or ‘so-called Paleozoic eugeosynclinal
sedimentary rocks’ in Japan (Koike 1979; Isozaki
& Matsuda 1980; Tanaka 1980; Yao et al. 1980,
Yamato-Omine Research Group 198 1; Nakaseko et
al. 1982); and (ii) the acceptance of a subduction-
accretion concept (Kanmera 1980; Taira et al.
1983) which was constructed mainly through deep-
sea drilling and regional seismic profiling on mod-
ern AC. As a result, most of the ‘so-called Paleo-
zoic eugeosynclinal rocks’ in Japan were revealed
to be Jurassic AC by the mid-1980s. Age of
accretion for each AC unit was precisely dated, and
this enabled mutual discrimination of neighboring
and similar-looking Paleozoic to Cenozoic AC in
Japan. This reform initiated a considerable redraft-
ing of the geotectonic history of the Islands in
terms of accretion tectonics, as well as their geo-
tectonic subdivision. Utility of microfossil dating
for ancient AC and the relevant results in geotec-
tonic subdivision of Japan up to the late 1980s are
summarized in Ichikawa et al. (1990).
On the other hand, concerning tectonic inter-
pretation, the ‘allochthonous or suspect terrane’
concept (Jones et al. 1977; Coney et al. 1980;
Howell 1985) invaded Japan in the early 1980s
and allowed many geologists in Japan to believe
the occurrence of exotic continental and/or oceanic
blocks and to use the term ‘terrane’ to describe
various orogenic units (Saito & Hashimoto 1982;
Mizutani 1987; Ichikawa et al. 1990). On the
geotectonic subdivision, the ‘terrane’-based under-
standing of the attitude of boundary faults
between the orogenic units is of note because most
of the ‘terrane boundaries’ were regarded as
vertical faults of a strike-slip nature related to
‘terrane dispersion’ (Taira et al. 1983). The nappe
tectonics, on the contrary, was revived also in
the early 198Os, and its significance with subhori-
zontal boundary faults was emphasized by French
and domestic geologists (Hara et al. 1977;
Yamato-Omine Research Group 1981; Faure 1985;
Charvet et al. 1985; Hayasaka 1987). French
geologists, in particular, interpreted the nappe-
related subhorizontal structure as a result of an
ancient continent-microcontinent collision, how-
ever, evidence for the putative collided micro- continent per se was not persuasive. Independent
from these works, relative plate motions were
partly reconstructed for the late Mesozoic and
Cenozoic Pacific region by Engebretson et al.
(1985), and the correlation between the plate
interactions in East Asia and orogenic events re-
corded in Japan was first discussed by Maruyama
and Sen0 (1986). Owing to the mixed effects of
these various interpretations and existence of prob-
lematic ‘gray zones’ mentioned below, understand-
ing of the 3D structure and tectonic evolution of
the Japanese orogen was in a state of confusion in
the 1980s.
The second wave in the early 1990s that pro-
vided the final tool for redrawing the subdivision of
the Islands came in the form of success in chrono-
metric dating of weakly metamorphosed AC. AC
well-recrystallized by regional metamorphism, such
as the Sanbagawa blueschists, had already been
dated by radiometric methods in the 1970s, how-
ever, less recrystallized metamorphic AC had not
been dated at all owing to difficulty in mineral
separation techniques. Similarly, while microfossil
analysis is powerful in dating non- to weakly
metamorphosed AC, those metamorphosed to the
greenschists facies were mostly left untouched
owing to difficulty in microfossil extraction. Such a
dilemma had left a considerable number of undated
weakly metamorphosed AC units, the ‘gray zones’,
in Japan even after the 198Os, and this delayed the
completion of the geotectonic subdivision of the
Japanese Islands. In the late 1980s, an advanced
technique of fine-grained mineral separation was
developed (Nishimura et al. 1989) that solved the
‘gray zone’ problem (Isozaki et al. 1990a, 1992;
Isozaki & Itaya 1991; Suzuki et al. 1990; Takami
et al. 1990; Kawato et al. 1991). Not only did this
permit the dating of the greenschist facies AC
rocks, but it allowed subdivision of high-PIT re-
gional metamorphosed units on the same basis as
non- to weakly metamorphosed ones (Isozaki &
Maruyama 1991). By 1993, most of the important
geotectonic boundaries were clearly defined by age
difference and were re-examined in the field. This
clarified that these boundary faults, that is, the
nappes, are essentially subhorizontal, except for
several vertical ones of secondary origin. In other
words, there is no deep crust-cutting vertical fault
zones nor a collage of exotic terranes in Japan. The
predominance in subhorizontal structures of the
Japanese Islands are currently accepted by many,
and are utilized, for example, in the latest version
of a geologic map of the islands issued by the
Geological Survey of Japan (1992).
14401738, 1996, 3, Downloaded from https://onlinelibrary.wiley.com/doi/10.1111/j.1440-1738.1996.tb00033.x by Ohio State University University Libraries, Wiley Online Library on [12/03/2025]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
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Island Arc - September 1996 - Isozaki - Anatomy and genesis of a subduction_related orogen A new view of geotectonic.txt
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CRETACEOUS RESEARCH ELSEVIER Cretaceous Research 25 (2004) 365390 www.elsevier.com/locate/CretRes Geologyandstratigraphyofforearcbasinsedimentsin Hokkaido, Japan: Cretaceous environmental events on the north-westPacific margin Reishi Takashimaa, *,Fumihisa Kawabeb,Hiroshi Nishi, Kazuyoshi Moriya, Ryoji Wani,HisaoAndo° Department of Earth and Planetary Science, Graduate School of Science,Hokkaido University,N10w8,Kita-ku,Sapporo 060-0810,Japan Institute of Natural History,3-14-24 Takada,Toshima-ku, Tokyo 171-0033,Japan Department of Earth Science,Graduate School of Social and Cultural Studies,Kyushu University，4-2-1 Ropponmatsu, Chuo-ku,Fukuoka 810-8560, Japan Deparment of Geology，NationalScienceMuseum323-1Hyakunincho,Shinjuku-k,Tokyo169-0073Japan °Department of Environmental Sciences,Faculty of Science, Ibaraki University, 2-1-1,Bunkyo,Mito 310-8512, Japan Received 13 August 2003; accepted in revised form 3 February 2004 Abstract Litho-,bio-, and chemostratigraphy of the Cretaceous forearc basin sediments exposed in Hokkaido, northern Japan allow a synthesis of the faunal, sedimentological, and environmental history of the north-west Pacific margin. Although the succession, named the Yezo Group, has yielded an abundant record of mid- to late Cretaceous invertebrates,monotonous lithologies of sandstone and mudstone, showing occasional lateral facies changes,have caused confusion regarding the lithostratigraphic nomenclature. Based on our wide areal mapping of the sequence, and analysis of litho- and biofacies, a new lithostratigraphic scheme for the Yezo Group is proposed. In ascending order, the scheme is as follows: the Soashibetsugawa Formation (Lower Aptian mudstone unit); the Shuparogawa Formation (Lower Aptian-lower Upper Albian sandstone-dominant turbidite unit); the Maruyama Formation (lower Upper Albian tuffaceous sandstone unit); the Hikagenosawa Formation (Upper Albian—Middle Cenomanian mudstone-dominant unit); the Saku Formation (Middle Cenomanian—Upper Turonian sandstone-common turbidite unit); the Kashima Formation (Upper TuronianLower Campanian mudstone-dominant unit); and the Hakobuchi Formation (Lower CampanianPaleocene shallow-marine sandstone-conglomerate unit). In addition, we designate two further lithostrati- graphic units, the Mikasa Formation (Upper AlbianTuronian shallow-marine sandstone-dominated unit) and the Haborogawa Formation (Middle TuronianCampanian shelf mudstone/sandstone unit), which correspond in age to the shallower facies of the Saku and Kashima formations, respectively. Despite a lack of so-called *black shales", because of siliciclastic dilution, our stratigraphic integration has revealed the horizons of oceanic anoxic events (OAEs) in the Yezo Group. The OAEla horizon in the Soashibetsugawa Formation is characterized by a lack of foraminifers, macrofossils and bioturbation, and a prominent positive excursion of 8Corg. A significant hiatus during the late Aptian and early Albian removed the OAE1b horizon. The OAElc horizon in the Maruyama Formation shows a distinct negative excursion of 'Corg with a concomitant high productivity of radiolarians. The OAE1d horizon in the middle part of the Hikagenosawa Formation consists of weakly laminated, pyrite-rich mudstone. Planktonic and calcareous benthic foraminifers are absent, whereas radiolarians are abundant above the OAE1d horizon. The mid-Cenomanian event (MCE) horizon is identified at the top of the Hikagenosawa Formation. Stepwise extinction of calcareous benthic foraminifers and a decrease in radiolarian diversity become apparent above the MCE horizon. In the study area, the OAE2 horizon has been well documented, and is placed in the middle part of the Saku Formation. 2004 Elsevier Ltd. All rights reserved. Keywords: Cretaceous; Stratigraphy; Yezo Group; OAEs; Pacific; Forearc basin * Corresponding author. E-mail address: rtaka@nature.sci.hokudai.ac.jp (R. Takashima) 0195-6671/S - see front matter 2004 Elsevier Ltd. All rights reserved. doi:10.1016/j.cretres.2004.02.004 366 R.Takashima et al./Cretaceous Research 25（2004）365-390 1. Introduction same lithostratigraphic unit. We, therefore, propose Knowledge of Cretaceous ocean-climate systems here a new, synthesized stratigraphic framework, in- provides important information to help understanding cluding macro- and microfossil biostratigraphy, based a future warm greenhouse world. In particular, the mid- on extensive mapping of the entire area in the central Cretaceous warming around 120-90 Ma, enhanced by region of Hokkaido. Moreover, we organize the strati- a huge production of oceanic crust and plateaus in the graphic and geographic distribution of OAE horizons Pacific and Indian oceans, brought higher levels of recorded in the Yezo Group using the new standard greenhouse gases and high poleward heat transportation stratigraphy. (e.g, Larson, 1991a,b). This global warming induced a unique deep-ocean circulation and expansion of anoxic conditions, known as oceanic anoxic events (OAEs) that resulted in large turnovers of marine biota (Leckie 2. Geological setting et al., 2002). Many palaeontological and geochemical studies of OAEs have been conducted in land sections The Jurassic to Palaeogene sequences exposed in the across the Tethyan region (e.g., Erbacher et al., 1996), western and central parts of Hokkaido, northern Japan, on deep-sea cores from the Atlantic (e.g., Bralower are basically divided into three N-S trending tectono- et al., 1994; Wilson and Norris, 2001), the Antarctic stratigraphic divisions called the Oshima (Cretaceous (e.g., Bralower et al., 1993), and from equatorial Pacific volcanic arc on a Jurassic accretionary complex), the regions (e.g., Sliter, 1989), whereas there is little infor- Sorachi-Yezo (Cretaceous-Paleocene forearc basin mation about Cretaceous environmental events in the and accretionary complex), and the Hidaka (lower North Pacific. Palaeogene accretionary complex) belts from west to The Cretaceous Yezo Group, exposed in central east (Kiminami et al., 1986; Ueda et al., 2000) (Fig. 1). Hokkaido, northern Japan, was probably deposited at These were formed along a westward-dipping sub- about 35-45°N (Hoshi and Takashima, 1999; Kodama duction of the Izanagi-Kula plates under the eastern et al., 2002) along a westward subduction margin in the margin of the Asian continent (e.g., Okada, 1974). north-eastern Asian continent during the Cretaceous. The Sorachi-Yezo Belt consists of a coherent This group consists of a 10,000-m-thick forearc sedi- succession, from the Horokanai Ophiolite through the ss n sss ns Sorachi Group to the Yezo Group (Fig. 2), and the subordinate conglomerates. This forearc basin is called accretionary complexes of the Idonnapu and Kamuiko- the Yezo Basin (Okada, 1983), and it extends from tan zones (Ueda et al., 2000; Fig. 1). The Horokanai offshore of north Honshu, through Hokkaido,to Ophiolite and the lower part of the Sorachi Group Sakhalin Island, Russia (Fig. 1). As the rocks contain represent a piece of basaltic oceanic crust (Ishizuka, abundant, well-preserved macro- and microfossils, many 1981; Takashima et al., 2002a), while the upper part of biostratigraphic schemes have been established (Matsu- the Sorachi Group is represented by subaqueous calc- moto, 1942, 1977; Tanaka, 1963; Obata and Futakami, alkaline and alkaline volcano-sedimentary sequences, 1975; Futakami, 1982; Taketani, 1982; Maeda, 1986; suggesting an oceanic island arc setting (Girard et al., Toshimitsu and Maiya, 1986; Motoyama et al., 1991; 1991; Nida, 1992; Takashima et al., 2002b). The Yezo Kawabe et al.,1996, 2003; Takashima et al., 1997; Group conformably overlies the Sorachi Group and Toshimitsu et al., 1998; Kawabe, 2000; Wani and comprises very thick sandstone and mudstone sequen- Hirano, 2000; Moriya and Hirano, 2001; Ando et al., ces. The sand clastics of this sequence were derived from 2001; Nishi et al., 2003). In the last decade, carbon Cretaceous granitic rocks and Jurassic accretionary isotope excursions of organic materials have also been complexes of the Oshima Belt, which represents a reported across the Cenomanian/Turonian boundary contemporaneous continental arc setting (Kito et al., (OAE2) and from the upper Lower Aptian (Hasegawa 1986). Although a part of the Hakobuchi Formation and Saito, 1993; Hasegawa, 1995, 1997; Ando et al., includes the upper Paleocene in the Oyubari and 2002,2003). Moreover, the thermal structure of the Nakatonbetsu areas (Ando et al., 2001; Ando, 2003), north-western Pacific of the Asian continental margin the geological age of this formation ranges mostly from has recently been investigated, based on oxygen isotopic early Aptian to early Maastrichtian. Many hiatuses exist analyses (Moriya et al., 2003). These studies provide between the Campanian and Maastrichtian. The sedi- important information of Cretaceous environmental mentary environment of this group shows an eastward- change and faunal turnover related to global events in deepening facies trend, from fluvial to continental slope the mid-latitude north-west Pacific. (Fig. 2). The Yezo Group is unconformably capped by However, the nomenclature of the lithostratigraphic late Eocene, non-marine and shallow-marine sediments divisions in the Yezo Group are complicated and differ of the Ishikari and Poronai groups, or by younger between the areas investigated within Hokkaido, and Neogene deposits. R.Takashimaetal./CretaceousResearch25（2004）365-390 367 SAKHALIN ISLA 35 30° Oshima Belt 25° Belt Hidaka Belt Teshionaka- Nakatonbetsuarea gawaarea OKHOTSK SEA Tomamae area (Fig. 5) 44°N Oyubari&Mikasa JAPANSEA area(Fig. 4 42N 42N Legend Cretaceous-Paleoceneforearcbasin sediments Palaeozoic to Lower Cretaceous O shelf to continental sediments C Upper Cretaceous to Eocene W EAN accretionary complexes Jurassic accretionary complexes UpperJurassictoLowerCretaceous volcano-sedimentarysuccessions Cretaceous volcano-sedimentary successions Cretaceous granitic rocks High-pressure-typemetamorphic rocks Ultramaficrocks Fig. 1. Simple geological map showing the distribution of the Mesozoic formations of north-east Japan and southern Sakhalin. Dashed-lined frames show the study areas (Figs. 4, 5). The central region (Fig. 4) consists of the south-eastern Oyubari area and the north-western Mikasa area. 368 R.Takashima et al./Cretaceous Research 25（2004）365-390 2.0 Lithology Planktonic 1.19 Age foraminiferal Environment Mikasa Tomamae,'Oyubari Mikasa Tomamae Oyubari zone area area area area area area CaO 172 Sst, coaly mdst and coal . Fluvial Group SS 5S Fluvial-inner Hakobuchi Fm Bioturbated sandy slst-sst,congl,HT sst Globotruncana arca shelf (<450m) and coaly mdst with felsic tuff Contusotrunca SS fornicata Tsukimi 56.81 Outer-inner ntor Sst Mbr. TCS sst Marginotruncana Strongly shelf sinuosa bioturbated 10.3 acian Haborogawa mdst-sandy mdst mdst Fm (1750-2250m） Felsic volcaniclastic sst (turbigite) and tuff 15.31 55.47 Turonian Sst>mdst Sst<mdst Bioturbatedmuddy sandstone 0.55 Hakkin helvetica Muddy I_Sst Mbr. Alternating beds of Saku Fm 0.44 (2300-1800m) Thick-bedded Rotalipora congl & sst , Sst<mdst Yezo Rotalipora Weaklybioturbated mdst globotruncanoides sandy mdst &massivemdst R.appenninid Kanajiri KY-3 Alte gbedsofsst（turbidite）&mdst "Barren interval" Sst>mds Rotalipora Weakly laminated and/or bioturbated mdst subticinensis- Rotalipora ticinens MaruyamaFm KY-2 Felsic volcaniclastic sst.(turbidite)and Biticinella （<500m） tuff beds withresedimented congl breggiensis jOkusakainosawa Ticinella Alternating beds of sst turbidite)& mdst Sst&MdstMbr Sst<mdst Sst>mdst primula Kirigishiyama Sst<mdst Olistostrome occasionally containing Ist listoliths Olisto. Mbr Globigerinelloides Sst<mdst Sst>mdst Sst<mdst Refureppu spp. Alternating beds of sst (turbidite) & mdst Sst Mbr sp<ss oashibetsugawa 20.9 0.91 Sorachi 57.0 of felsic tuffbeds Group Mafic-intermediatevolcaniclasticsandlava Oceanic Horokanai Ophiolite Mafic-ultramafic igneous basement islandarc Fig. 2. Schematic diagram of the Yezo Group in the study areas. Note that the geological column shows eastward-deepening facies. 3. Stratigraphy of the Yezo Group units, with intercalations of six distinct stratigraphic key 63.35 Lithostratigraphic nomenclature for the Yezo Group definitions proposed in this paper are as follows, in as- is very confused. Various definitions have been proposed, 63.08 depending on the area and the researchers, because this tion (mudstone unit); the Shuparogawa Formation 19.39 (sandstone-dominant turbidite unit); the Maruyama sandstone and mudstone and their alternating beds, Formation (felsic tuff and tuffaceous sandstone unit); which occasionally exhibit lateral facies changes. The the Hikagenosawa Formation（ (mudstone-dominant criteria of lithological boundaries have been defined by unit); the Saku Formation (sandstone-common turbidite sandstone intercalations and/or detailed changes of the unit); the Kashima Formation (mudstone unit) and sandstone/mudstone ratios in alternating beds. However, the Hakobuchi Formation (shallow-marine sandstone our study has revealed that the Yezo Group throughout unit). Hokkaido is basically characterized by six alternations We also recognize the sandstone-dominant, outer- of mudstone-dominant units and sandstone-common shelf to shoreface Mikasa Formation and the overlying, R.Takashima et al./Cretaceous Research 25（2004）365-390 369 Table 1 List of the widely traceable key units of the Yezo Group Key units Thickness Lithofacies Horizons Age Picture KY-6 10-30 m Tuffaceous coarse Middle parts of the Kashima Latest Santonian Fig. 7N, 0 sandstone beds and Haborogawa formations (Tsukimi Sst Mbr) KY-5 10-20 m Greenish tuffaceous Lower parts of the Kashima Early Coniacian Fig. 7M coarse sandstone and Haborogawa formations beds with abundant I. uwajimensis KY-4 <300 m Greenish gray Middle part of the Early Turonian Fig. 71 tuffaceous muddy Saku Formation sandstone with abundant (Hakkin Muddy Sst Mbr) Planolites trace fossils and very thick felsic tuff beds KY-3 100-300 m Sandstone-conglomerate Lower middle part of the Latest Albian Fig. 7G beds Hikagenosawa Formation (Kanajiri Sst Mbr) KY-2 <82 m Felsic tuff and tuffaceous Maruyama Formation Early Late Albian Fig. 7E sandstone beds KY-1 <400 m Olistostrome or debris fow Middle part of the Late Aptian- Fig. 7C deposits occasionally containing Shuparogawa Formation Early Albian "Orbitolina limestone blocks (Kirigishi. Olistostrome Mbr) sandy mudstone-dominant outer-shelf Haborogawa Lithology. The formation consists of dark grey Formation. The former is exposed in the Mikasa area, parallel-laminated mudstone with many intercalations and the latter in the Tomamae and Mikasa areas (Figs. 2, of felsic tuff beds (Fig. 7A). The laminae are composed 4, 5). The geological ages of these formations correspond of felsic tuff and very fine-grained sand layers. Deep-sea to the Saku and Kashima formations exposed in the trace-fossils, such as Lorenzinia and Cosmorhaphe, are Oyubari area, respectively (Fig. 2). Additionally, in the found occasionally in the mudstone. The tuff beds are study areas we name six stratigraphic key units as KY-1 white, hard and generally 10-30 cm thick, though some to KY-6 consisting of olistostrome (KY-1), tuffaceous beds attain a thickness of 1-7 m. They contain fine- sandstone (KY-2, KY-4 to 6), and sandstone-conglom- grained, bubble-wall glass shards with minor amounts of erate units (KY-3) (Fig. 2, Table 1). idiomorphic feldspar and biotite. Thickness and distribution. The formation is 450- 3.1. Soashibetsugawa Formation (redefined) 700 m thick in the eastern part of the Oyubari area, and Definition. This formation is characterized by a pre- not exposed in the Mikasa and Tomamae areas. dominance of dark grey siliceous mudstone, and cor- Fossils and age. No macrofossils have been found in responds to the Soashibetsugawa Mudstone Member of the Shuparogawa Formation of Takashima et al. (2001) this formation. Radiolarians occur abundantly through- (Fig. 3). Takashima et al. (2001) defined the member as out the sequence of this formation, but there are no age-diagnostic species. Planktonic foraminifers appear the basal unit of the Yezo Group because of the lithofacies shift from in situ volcanic and volcaniclastic rarely in the upper part (Fig. 8). The early Aptian spe- cies, Leupoldina cabri (Sigal), occurs from the upper- rocks of the Sorachi Group to terrigenous, dark grey mudstone at the base of the member. As the member is most part of the formation (Saito and Ando, 2000; widely traceable and clearly distinguishable from other Takashima et al., 2001; Nishi et al., 2003). Agglutinated members (sandstone-dominant facies) of the Shuparo- benthic foraminifers, such as Bathysiphon, also occur gawa Formation of Takashima et al. (2001), we redefine occasionally. the member as the basal formation of the Yezo Group. Depositional environment. Microfossil assemblages Stratigraphic relationship. This formation conform- and deep-sea trace fossils suggest an abyssal environment. ably overlies the Shirikishimanaigawa Formation of the Sorachi Group. 3.2. Shuparogawa Formation (redefined) Type section. The Soashibetsu River section, northern Definition. This formation, corresponding to the Oyubari area (Figs. 4, 6; section 11). Tomitoi and Shuparogawa formations of Motoyama 370 R.Takashimaetal./CretaceousResearch25（2004）365-390 Mikasa area Tomamae area Oyubari area This paper Nishida et al. Wani & Hirano Matsumoto Motoyama et al. Ando (1990b) This paper Takashima et al This paper (1997) (2000)* (1942) (1991)** (2001)* Hakobuchi Hakobuchi Hakobuchi Hakobuchi Hakobuchi Group Fm Fm Group Group Fm Tsukimi Upper Haboro- Upper IIld Upper gawaFm Haborogawa Ammonite Illc Kashima Kashima Yezo Group MiddleHaboro- Fm Group IIb Fm Fm Fm gawaFm Illa PM Shirogane Fm Twb TwC Mikasa Mikasa Fm Shirochi Fm Hakkin Muddyi Fn Hakkin Muddy dng Takinosawa Fm Twa My7-8 _Sst Mbr_- Middle lk-g Fm Mbrs Ammonite Tenkaritoge Hikagenosawa Group Main G lle G Hikagenosawa Md Part 6 Kanajiri Fm Ild My1-6 Fm Hikagen Kanajiri Fm Fm Mbrs Ilc Fm Fm Sst Mbr_i MI Ilb Takimibashi Fm MaruyamaFm MaruyamaFm lla Maruyama Fm MaruyamaFm MaruyamaFm Shuparogawa jOkusakai.Sst Okusakai. Sst &mdst Mbr Fm &mdst Mbr Lower Shup Olist.Mbr Ic Fm Ammonite Oolist.Mbr Group b Refureppu Refureppu SstMbp la Tomitoi Fm Sst Mbr Soashibetsu chiGro Sorachi Group Onisashi Group Sorachi Group Soashibetsugawa (Uppermost part) (Uppermost part) Mbr ed in Nishi et al.(2003) va (1994),Hasegawa (1995) and Hasegawa (1997) Fig. 3. Comparison of lithostratigraphic subdivisions proposed by several workers in the study areas. et al． (1991),is distinguished from the mudstone- ratio in the lower part is about 5/1. Sandstones in the dominant Soashibetsugawa Formation by the onset of upper part are less than 5 cm thick, with Te-e divisions, common to frequent intercalations of sandstone beds. and the S/M ratio is 1/2 to 1/5. Mudstones in this member are uniformly less than 10 cm thick, dark grey Type section. Along the Shuparo River section, the and weakly bioturbated. In the northern Oyubari area, central Oyubari area (Figs. 4, 6; section 12). this member is dominated by sandstone throughout (Fig. 6; section 11). Stratigraphic relationships. The Shuparogawa For- The Kirigishiyama Olistostrome Member represents mation conformably overlies the Soashibetsugawa an olistostrome bed containing huge allochthonous Formation in the Oyubari area, while the basal part of blocks of massive sandstone (<40 m thick), alternating this formation is not exposed in the Tomamae and beds of sandstone and mudstone (<20 m thick) and Mikasa areas because of poor exposure. limestone (<60 m thick) in a muddy matrix (Fig. 7C). Limestone olistoliths consist of corals, large foraminifers Lithology. The Shuparogawa Formation is mainly (Orbitolina), rudists and ooids, contaminated with composed of alternating beds of turbiditic sandstone pebbles of chert, granite and sandstone. The width and and dark grey mudstone. This formation incorporates thickness of limestone olistoliths is greatest in the a thick olistostrome bed (Kirigishiyama Olistostrome northern Oyubari area where a 60-m-thick, slab-shaped Member) in the middle portion, and is subdivided block is continuously exposed for a distance of 3 km in into three members as follows: the Refureppu Sand- a N-S direction (Fig. 4, around Mt. Kirigishiyama). The stone, Kirigishiyama Olistostrome, and Okusakaino- limestone olistoliths thin out in the Tomamae and sawa Sandstone and Mudstone members, in ascending southern Oyubari areas. However, the olistostrome bed order (Figs. 2, 6). is widely traceable throughout the study area as a key The Refureppu Sandstone Member consists of a very unit (KY-1) (Fig. 6). thick- to medium-bedded, sandstone-dominant sequence The OkusakainosawaS Sandstone and Mudstone in the lower part (Fig. 7B) and thin to very thin-bedded, Member basically comprises mudstone-dominant alter- mudstone-dominant alternating beds in the upper part. nating beds, and varying thicknesses of sandstone, from It contains slump beds locally. Lower sandstones are 5 to 50 cm depending on the area (Fig. 7D). Sandstones stratified, well sorted, usually ranging from 10 to 50 cm are turbiditic, exhibiting Tb-e sequences with sole marks thick, and occasionally more than 1 m thick. They are and numerous plant fragments, while thick sandstones turbidites showing the S3—Tb-e divisions of Lowe (1982), show the S1-3 divisions of Lowe (1982). In the Oyubari with many sole marks. The sandstone/mudstone (S/M) area, intercalations of sandstone become common to R.Takashima et al./Cretaceous Research 25（2004） 365-390 371 abundant in the middle part of this member (Fig. 6; 1.41 section 11). Mudstones are bioturbated with moderate 4.23 Definition. This formation is defined by an assem- blage of hard felsic volcaniclastic sandstones, tuffs and Thickness. The formation is approximately 1500 m associated conglomerates (Motoyama et al.，1991) thick in the type section. However, it varies in thick- (Fig. 3). It also yields an excellent stratigraphic marker 2.42 (KY-2), which extends throughout Hokkaido (Table 1). Fossils and age. Mudstones of the Refureppu Sand- Type section. Along the Shuparo River section, cen- stone Member yield a few foraminifers and common 0.44 radiolarians, but no macrofossils (Fig. 8). Planktonic foraminifers include the Upper Aptian species, Globiger- 1.45 inelloides barri (Bolli, Loeblich and Tappan) and G. 44.9 duboisi (Chevalier). Agglutinated and calcareous benthic foraminifers also occur occasionally. The Okusakaino- Lithology. This formation is composed of tuffaceous sawa Sandstone and Mudstone Member contains abun- sandstone beds (Fig. 7E), locally accompanied by dant radiolarians and planktonic foraminifers, with a conglomerates at the base. The tuffaceous sandstone subordinate proportion of common agglutinated and beds are 0.3-2 m thick, white to pale brown, siliceous calcareous benthic foraminifers. The Upper Albian 22.5 ammonoid, Mortoniceras cf. geometricum Spath, occurs graded to parallel-laminated sequences (Tb-e) and basal in the upper part of this member (Kawabe, 2000). This rip-up mudstone clasts. Microscopically, the sandstones member also includes the transitional interval from the are composed of platy glass shards, subordinate idio- Ticinella primula planktonic foraminifera (pf) Zone to morphic plagioclase, biotite, quartz and hornblende. The the lower part of the Biticinella breggiensis pf Zone, tuffaceous unit varies from 4 to 84 m thick throughout assigned to the Lower-lower Upper Albian (Nishi et al. the study area. 2003) (Fig. 8). A significant hiatus is present from The conglomerates accompanying the base of this the Planomalina cheniourensis toHedbergella planispira formation are usually less than 2 m thick. However, in zones of Hardenbol et al. (1998), approximately 115.5- NazO 108.21 Ma (7 myr). It might have been eroded during thickness of the conglomeratic beds exceptionally attains deposition of the Kirigishiyama Olistostrome Member. as much as 900 m (Fig. 6; section 10). The conglomerates The limestone olistoliths yield larger foraminifers of are thick-bedded, poorly sorted, and clast-supported, orbitolinids, e.g., Orbitolina lenticularis (Blumenbach), structureless or exhibiting R3—S1 divisions of Lowe 8.62 2.84 (Matsumaru, 1971). pebbles-boulders of rhyolite, with subordinate, well- rounded granules to pebbles of mudstone, chert and Depositional environment. Lithology and benthic 32.5 foraminiferal data suggest that this formation was the Oyubari area they are composed of granules to deposited on the continental slope (e.g., Motoyama pebbles of chert, sandstone, siliceous mudstone and et al., 1991). Benthic foraminifers of the Refureppu limestone. 28.7 and Mudstone Member, consisting of species of Bathy- Thickness. This formation attains thicknesses of siphon, Gaudryina, Gyroidinoides, Gavelinella and Nodo- 1 saria, indicate the upper bathyal zone or deeper (e.g., Oyubari area. The horizon and thickness of this for- Sliter and Baker, 1972). However, the limestones of mation are disputable in the Mikasa area because of the Kirigishiyama Olistostrome Member are considered poor exposure. to have been formed on a carbonate platform, such as a rimmed shelf along the Asian continental margin (Sano, Fossils and age. Although no planktonic foraminifers 1995). The olistostrome indicates the collapse of the and macrofossils have been obtained from this forma- shallow carbonate platform and mixing with the bathyal tion, radiolarians (spumellarians） occur abundantly mudstone around the Aptian/Albian boundary. (Fig. 8). Fig. 4. Geological map and structural profile section of the Oyubari and Mikasa areas, central Hokkaido Island region. The south-eastern Oyubari and north-western Mikasa areas are bounded by the Katsurazawa Thrust Fault. Although most areas in this map are based on our original data, the southern Mikasa area (the Manji and Hatonosu domes) and northeastern Oyubari area (around the Mt. Yubaeyama) are modified from Tanaka (1970), Obata and Futakami (1975), Futakami (1982) and Hashimoto (1953), respectively. 372 R. Takashima et al./Cretaceous Research 25(2004） 365-390 ranoCity YezoGroup R.Takashima et al./Cretaceous Research 25(2004） 365-390 373 YezoGroup 374 1031m YezoGroup beushinaiVill. BD √B' Mt.Sar 10km OTappuVill. 2000 48 4B' 1000 Fig. 5. Geological map and structural profile section of the Tomamae area; see Fig. 4 for legend. R.Takashima et al./Cretaceous Research 25（2004） 365-390 375 Depositional environment. Abundant volcaniclastic igneous rocks. The horizon and thickness of this materials in this formation suggest that huge, felsic member are disputable in the Mikasa area because of volcanic eruptions episodically occurred along the poor exposure and scarce fossil data. western circum-Pacific/Asian continental margin. Main transporting channels for the volcanic sediments are Thickness and distribution. 1900-2600 m thick in the inferred to have been located in the Tomamae area. Oyubari area, and 1900-2000 m thick in the Tomamae area. The thickness in the Mikasa area is uncertain 3.4.Hikagenosawa Formation because of poor exposure. Definition. This formation is defined by the pre- Fossils and age. The occurrence of macrofossils dominance of dark grey mudstone (Motoyama et al., (ammonoids and inoceramids) becomes common above 1991; Fig. 3), although there are intercalations of a thin the lower part of this formation (Fig. 8). The late sandstone-dominant unit (Kanajiri Sandstone Member) Albian-middle Cenomanian ammonoids Mortoniceras in the lower middle part. rostratum (Sowerby), Mariella bergeri (Brongniart), Mantelliceras saxbii (Sharpe) and Cunningtoniceras Type section. Along the Hikageno-sawa Valley sec- cunningtoni (Sharpe) occur in the formation. Micro- tion, the central Oyubari area (Fig. 4). fossils are abundant throughout the sequence. Excep- tionally, planktonic foraminifers and calcareous benthic Stratigraphic relationship. The formation overlies, foraminifers are lacking around the Kanajiri Sandstone conformably, the Maruyama Formation. Member, whereas radiolarians become very abundant and agglutinated benthic foraminifers are common in Lithology. The dominant lithology of the Hikageno- this interval (Fig. 8). Planktonic foraminifers indicative sawa Formation is dark grey mudstone (Figs. 6, 7F). A of the Biticinella breggiensis to Rotalipora cushmani pf thin unit of alternating beds of sandstone/conglomerate zones, for example, Biticinella breggiensis (Gandolfi), Ticinella subticinensis (Gandolfi), Rotalipora appenninica and mudstone (Kanajiri Sandstone Member) is in- tercalated in the lower middle part of this formation. (Renz), R. globotruncanoides Sigal, and R. cushmani This unit is used as a stratigraphic marker, the KY-3 (Morrow), occur, indicating a Late Albian-Late Cen- (Fig. 2, Table 1). omanian age (Nishi et al., 2003; Fig. 8). The lower part of this formation is characterized by weakly-laminated mudstone. The intercalations of thin- Depositional environment. Benthic foraminiferal as- bedded felsic tuffs and distal turbiditic sandstones semblages suggest that this formation was deposited in showing Tc-e divisions are common in the Tomamae the lower part of the upper bathyal zone under relatively and Mikasa areas. In the northern Tomamae area, oxygenated conditions (Motoyama et al., 1991; Kaiho thick- to very thick-bedded sandstones with S3—Tb-e are, et al., 1993). exceptionally, also intercalated (Fig. 6; section 9). On the other hand, rare sandstones are intercalated in the 3.5.Saku Formation Oyubari area. Mudstones in the upper part of this formation are moderately bioturbated, and occasionally Definition. This formation is defined by the frequent weakly laminated. intercalations of turbiditic sandstone beds (Matsumoto, The Kanajiri Sandstone Member is composed of 1942). thick- to very thick-bedded sandstone with very thick- bedded conglomerates in the Tomamae area, whereas in Type section. The Abeshinai River section, Teshio- the Oyubari area it comprises thin- to medium-bedded, nakagawa area, northern Hokkaido (Fig. 1). The mudstone-dominant alternating beds of sandstone and lithology and distribution of this formation were well mudstone (Fig. 7G). Sandstones are proximal (S3-Tb-e) documented by Matsumoto (1942) and Matsumoto and in the former area and distal (Ta-e) in the latter. The Okada (1973). conglomerates are well-rounded pebbles to granules, exhibiting R3-S1 divisions. Their compositions are Stratigraphic relationship. The formation overlies con- mostly chert with minor sandstone, mudstone and formably the Hikagenosawa Formation. Fig. 6. Correlation of selected sections from the study areas. 1-5, Mikasa area. 1, Shikoro-zawa Valley; 2, Ponbetsu River; 3, Yonno-sawa Valley; 4, Ganseki-zawa Valley-Nanashi-zawa Valley; 5, Tsukimi-sawa Valley. 6—10, Tomamae area. 6, Sankebetsu River; 7, Nakafutamata River; 8, Shumarinai River; 9, Sounnai River-Kotanbetsu River; 10, Gosen River-Kanajiri-zawa Valley-Akanosawa Vally. 11—17, Oyubari area. 11, e zawa Valley. Each section is cited in Figs. 4 and 5. Note that the precise thicknesses of the KY2 (sections 11-13) and KY-5,6 could not be depicted because of scale problems; see Table 1 for these. HakobuchiFm KY-5 HaborogawaFm HikagenosawaFm Tomamaearea Sections6-10,Figure 5) Mikasaarea (Sections1-5 Oyubariare Figure (Sections 11-17, Figure4) 1000m 500 YezoGroup Oyubariarea 。。 O C 150 7 O ：16 HakobuchiFm 1000m 11 13 KashimaFm 500 14 -0 KY-5 Legend 12 cross-stratification Troughcross-stratification Carbonaceousmudstone Laminatedmudstone Weaklylaminatedmudstone Massive mudstone Weaklylamina Weaklylaminatedmuddysandston Massivemuddysandstone SakuFm Fine-grained sandstone Coarse-grained sandstone HakkinMuddy Conglome obble) SstMbr KY-4 Resedimented (>cobble) conglomerate (granule-pebble) (>1m) ebed (turbidite) Felsicvolcaniclastic Hikagenosawa Fm (>1m） (3-30cm) Kanajiri SstMbr KY-3 Shellbed MaruyamaFm awaSst&MdstMbr Alternatingbedsof maOlistMbr KY-1 natingbedsofsst&mdst (foldedsstdominantalter- nating bedsofsst&mdst Fault SoashibetsugawaFm 1 378 R.Takashima et al./Cretaceous Research 25（2004）365-390 Lithology. This formation comprises alternating beds The sandstones are very thick-/thick-bedded, showing of turbiditic sandstone and mudstone, and intercalates S1—Tb-e divisions. The second and fourth cycles show the with a unit characterized by a predominance of greenish facies changes from very thick-/thick-bedded, sandstone- grey muddy sandstones (Hakkin Muddy Sandstone dominant alternating beds to thin-bedded mudstone- Member; KY-4) in the middle part (Fig. 2, Table 1). dominant ones. The sandstones are proximal turbidites, Sandstones in the lower part (below the KY-4) are exhibiting S1—Tb-e divisions. less than 40 cm thick and exhibit typical turbiditic sequences, having Tb-e divisions. Mudstones are dark Thickness and distribution. 2300 m in the Oyubari grey-coloured and moderately bioturbated. The S/M area, 1800-2000 m in the Tomamae area. The upper ratios change from 1/10 to 1/1 (Fig. 7H). part of this formation is truncated by the N—S trending A very characteristic lithofacies, greenish grey muddy thrust fault (Kashima Thrust Fault) in the northern sandstone with abundant trace fossils of Planolites Oyubari area (Fig. 4). (Fig. 7I), is observable in the middle part of this formation. This greenish unit extends from the Oyubari Fossils and age. Except for the Hakkin Muddy area to the central Tomamae area, about 200 km, and Sandstone Member the formation contains abundant is defined as the KY-4 (Hakkin Muddy Sandstone macrofossils (Fig. 8). The Cenomanian ammonoids, Member). The member incorporates medium-/thick- Calycoceras spp., occur in the lower part of the for- bedded sandstones, with S3-Tb-e divisions, and frequent mation. The blackish-grey mudstone intercalated in the felsic tuff beds. The tuff beds vary in thickness from 0.02 basal part of the Hakkin Muddy Sandstone Member is to 2 m, and are altered to bentonites. Finely laminated, correlated with the Cenomanian/Turonian boundary, pyrite-rich, dark grey-greyish black mudstones are based on the results of carbon isotope, mega- and intercalated just below or in the lower part of this microfossil stratigraphy (e.g., Hasegawa and Saito, member in the Oyubari and Tomamae areas. 1993; Hirano, 1995). Although macrofossils are few in The upper part of the formation begins with dark the Hakkin Muddy Sandstone Member, Lower Turo- grey mudstone and changes to alternating beds of nian indicators, such as Pseudaspidoceras fexuosus sandstone and mudstone.Sandstonesare thin-to (Powell) and vascoceratids, occur. The upper sequence medium-bedded turbidites with Tb-e sequences, and the is marked by an abundant occurrence of heteromorphic S/M ratios are not high, 1/10 to 1/5. Slump beds and ammonoids, including Nipponites, Eubostrichoceras and sandstone dykes are common in the southern Oyubari scaphitids, with minor Middle Turonian indicators, such area. Sandstones are occasionally intensively biotur- as Romaniceras spp. (ammonoid) and Inoceramus bated and bedding planes of some beds are destroyed. hobetsensis Nagao and Matsumoto (bivalve). Micro- Mudstones and sandy mudstones are dark grey to grey fossils commonly occur throughout the sequence; radio- and intensively bioturbated. larians are especially abundant in the Hakkin Muddy On the other hand, the Saku Formation, exposed Sandstone Member, and occupy 80-90% of the total. in the northern Tomamae area, is marked by four This formation ranges from the Rotalipora cushmani to thinning-upward sequences (Fig. 6; section 8). A channel- Helvetoglobotruncana helvetica pf zones, assigned to the filling conglomerate/sandstone unit constitutes the Upper Cenomanian-Middle Turonian (Nishi et al., first and third sequences. The conglomerates are clast- 2003; Fig. 8). supported, and 0.5—4 m thick, with disorganized erosional bases (Fig. 7J). Graded bedding, from well-rounded cob- Depositional environment. The benthic foraminif- ble to coarse sandstone (R2-3—S1 divisions), is common. eral assemblages indicate the lower part of upper Fig. 7. Photographs of the representative lithofacies of the Yezo Group. A, laminated mudstone with frequent intercalations of felsic tuff beds; Soashibetsugawa Formation, Soashibetsu River, northern Oyubari area. B, sandstone-dominant alternating beds of sandstone and mudstone; Shuparogawa Formation, Refureppu Sandstone Member, Soashibetsu River, northern Oyubari area. C, olistostrome containing large limestone blocks; Shuparogawa Formation, Kirigishiyama Olistostrome Member (KY-1), Okusakaino-sawa Valley, central Oyubari area. D, alternating beds of sandstone and mudstone; Okusakainosawa Sandstone and Mudstone Member, Shuparo River, central Oyubari area. E, felsic tuffaceous sandstones; Maruyama Formation (KY-2), Uenbetsu River, southern Tomamae area. F, weakly laminated mudstone; Hikagenosawa Formation, Tengu-sawa Vally, central Oyubari area. G, mudstone-dominant alternating beds of sandstone and mudstone; Hikagenosawa Formation, Kanajiri Sandstone Member (KY-3), Shuparo River, central Oyubari area. H, alternating beds of sandstone and mudstone; Saku Formation, Shumarinai River, northern Tomamae area. I, greenish grey muddy sandstone with abundant Planolites (KY-4), Saku Formation, KY-4, Hakkin-zawa Valley, southern Oyubari area. J, conglomerate bed with erosional base; Saku Formation, Shumarinai River, northern Tomamae area. K, sandstone with hummocky cross-stratification; Mikasa Formation, western Katsura-zawa Lake, southern Mikasa area. L, massive mudstone; Kashima Formation, Horokakuruki River, southern Oyubari area. M, Inoceramus uwajimensis shell bed (KY-5), Kashima Formation, Okusamata-zawa Valley, central Oyubari area. N, felsic volcaniclastic sandstone; Haborogawa Formation, Tsukimi Sandstone Member (KY-6), Ashibetsu River, eastern Mikasa area. O, coarsening upward sequences, from sandy mudstone to medium-grained sandstone; Haborogawa Formation, Nakanofutamata River, northern Tomamae area. P, sandstone with hummocky cross-stratification; Hakobuchi Formation, Hachigatsu-zawa Valley, eastern Mikasa area. R.Takashima et al./Cretaceous Research 25(2004） 365—390 379 erosionalbase 30cm Fig. 7 (continued) R.Takashima et al./Cretaceous Research 25（2004）365-390 381 Bioticevents Planktonic foraminifers Macrofossils Organiccarbon Oceanic Planktonic Litho- foraminiferal Stage 2.72 Composite 0.35 8.3 2.65 column events 2.87 units 3.64 diversity biohorizons Nishiet al. (2003) -23 -22 9.63 Fm nian 924 G. arca Tsukimi 0.3 iSst Mbr Santonian 1.45 C. fornicata 42.0 1.75 0.7 0.30 M.pseudolinneia nian H.helvetica 5.29 Hakkin 12.0 0.42 0.37 Cenomanian 2.92 1.84 Kanajiri !Sst Mbr .tici-R.subticinens 214 B.breggiensis 0.46 0.32 Lackof 0.52 bejaouensisZones 6.12 Globigerinelloides 65 G.ferre 7.9 2.28 L. cabri 296 ashibets 17.2 Pyrite-rich -4genera Fig. 8. Composite column of the Yezo Group, summarized lithology, macro-, microfossil and carbon isotope stratigraphy and microfossil biotic events. Note that five OAE horizons are identified within the Yezo Group. 14.3 Type section. North-west of Katsurazawa Lake, in 13.9 76.6 Lithology. The Mikasa Formation is characterized by 14.7 a predominance of sandstones with hummocky and 92.8 Definition. This formation is the contemporaneous, shell lags, and intensively bioturbated sandy mudstone shallower facies of the Saku Formation, and is defined 47.4 by the predominance of sandstones exhibiting hum- successions, assigned as third-order depositional sequen- mocky and trough cross-stratification (Matsumoto, ces (DS1, DS2 and DSs), each of which further includes 1951; Ando, 1990a). three fourth-order sequences (4th DS) (Ando, 1997), 382 R.Takashima et al./Cretaceous Research 25（2004）365-390 Within each DS, conglomerate and coarse-grained sand- been deposited in lower shoreface to outer shelf stone tend to thicken in the west, and interbedded sandy environments. mudstone becomes dominant and thick in the east. In the northwestern Mikasa area, the DS, begins with 3.7.KashimaFormation alternating beds of turbiditic sandstone and mudstone, and grades into HCS sandstone with bioturbated, very Definition. The formation is defined by the pre- fine-grained sandstone/sandy mudstone. A basal thick, dominance of dark grey, massive mudstone overlying cross-stratified conglomerate of the DS2 covers HCS the Saku Formation (Motoyama et al., 1991; Fig. 3). sandstones of DSi with a sharp erosional base (un- conformity) (Fig. 6; section 2). The DS2 comprises Type section. Kashima village in the southern a wide variety of lithofacies, such as cross-stratified Oyubari area (Fig. 4). conglomerate, carbonaceous mudstone, sandy mud- stone with oyster-shell beds, medium- to coarse-grained Stratigraphic relationship. This formation covers, con- sandstone with TCS, HCS fine sandstone and biotur- formably, the Saku Formation. bated sandy mudstone. The DS3 deposits are distributed in the southern Mikasa area, mainly comprising bio- Lithology. This formation consists mainly of dark turbated, sandy mudstone and HCS sandstones. The grey, bioturbated, massive mudstone (Fig. 7L), grading intercalations of HCS sandstone become frequent in into muddy sandstones in the uppermost part. The the upper part of each fourth DS. These DSs grade into mudstones are intercalated with two felsic volcaniclastic the Saku Formation northward and eastward. units, in the lower and middle parts, respectively. The lower volcaniclastic unit is 10-20 m thick, consisting of Thickness and distribution. This formation is exposed thick-bedded, coarse- to fine-grained, volcaniclastic only in the Mikasa area, forming the Sorachi-Ikush- sandstones with interbedded dark grey mudstones. The umbetsu Anticline and the Manji and Hatonosu domes. sandstone beds are turbiditic with Tp-e divisions, and less Its thickness ranges from 400 to 750 m, and it tends to than 2 m thick, rarely attaining 4 m thick in the northern thicken eastward. Oyubari area. The thickness of the mudstones is variable, depending on the area, but all are less than Fossils and age. Shallow-marine bivalves (including 2 m thick. A diagnostic feature of this unit is abundant Apiotrigonia, Glycymeris, Meekia, Pinna, Pseudoptera, occurrences of Inoceramus uwajimensis Yehara from the Pterotorigonia, Thetis, Yaadia), gastropods (e.g., Mar- interbedded mudstone (Fig. 7M). This unit is an ex- garites, Semisolarium) occur as storm-lag deposits within cellent stratigraphic marker (KY-5) throughout the HCS sandstones. Bioturbated, muddy sandstones to study areas (Table 1). sandy mudstones contain rare to common ammonoids Compared with the lower unit, the upper volcani- and inoceramids, but no microfossils. The Cenomanian clastic unit differs in lacking mudstone intercalations ammonoids, Mantelliceras sp., Calycoceras sp. and and macrofossils. Tuffaceous sandstones or tuffs are Desmoceras spp. have been reported from the DS1 thick-bedded (<1 m thick), white, coarse- to very fine- deposits (Kawabe, 2003). The Turonian species, Inocer- grained. Some are altered to bentonite. amus hobetsensis and Inoceramus teshioensis Nagao and Matsumoto, are abundant in bioturbated sandy Thickness. About 1670 m in the southern Oyubari mudstones of DS2 and DS3, respectively (Ando, area. The thickness of this formation in the northern 1990a,b). Oyubari area is difficult to estimate because of the monotonous lithology and complex geological structure, Depositional environment. Three DSs show conspicu- with isoclinal folding and thrust faulting. ous lateral and vertical facies changes, representing repetitive delta progradations in the western margin of Fossils and age. This formation contains abundant the Yezo Basin. Judging from the stacking patterns of calcareous concretions, including macrofossils such as their facies, third- and fourth-order DSs, the delta ammonoids and inoceramids. I. uwajimensis is particu- system is presumed to have shifted southward within the larly abundant in the KY-5 unit (Fig. 8). distribution area (Ando, 1997, 2003). The depositional Microfossils (foraminifers and radiolarians) are environment of the DSi deposits is interpreted as basin- plain on continental slope to lower shoreface through rare. The following four mid-latitude zones of plank- shelf. The DS2 represents non-marine to shallow-marine tonic foraminifers have, therefore, been proposed from environments, such as fuvial channel, back marsh, the Late Turonian to Campanian in the study area, flood plain and tidal flat in the lower part, followed by namely: the Marginotruncana pseudolinneana, M. sinu- deeper marine (shoreface to inner shelf, though partly osa, Contusotruncana fornicata, and Globotruncana arca outer shelf) in the upper. The DSs is thought to have pf zones (Fig. 8). These can be correlated directly with R.Takashima et al./Cretaceous Research 25（2004）365-390 383 the Tethyan planktonic foraminiferal zones (Nishi et al., Thickness. Approximately 1950 m in the Mikasa area, 2003). 1750-2250 m in the Tomamae area. Depositional environment. The benthic foraminiferal Fossils and age. This formation yields abundant, well- assemblages indicate the upper part of the upper bathyal preserved macrofossils from calcareous concretions in environment with medium- to relatively high-oxygen both the Tomamae and Mikasa areas (e.g., Futakami, levels (Kaiho et al., 1993). 1986; Maeda, 1986; Wani, 2001; Moriya et al., 2003). Several regional biozones, the I. uwajimensis-I. mihoensis, 3.8. Haborogawa Formation (new) I. amakusensis, and 1. japonicus inoceramid zones, have been proposed in the Tomamae area by Toshimitsu and Definition. The base of this formation is delineated by Maiya (1986). Microfossils are abundant throughout the bioturbated mudstone without sandstone intercalations, study areas. Although Tethyan planktonic foraminifers overlying the Saku Formation. While the lower part of are sporadic, the four mid-latitude zones, from the this formation is mainly composed of bioturbated Marginotruncana pseudolinneiana to the Globotruncana mudstone, the upper part is characterized by coarsening arca pf zones, can be identified in this formation. Its upward successions, from mudstone to muddy sand- geological age is considered to be Coniacian--Campanian stone and/or sandstone. (Moriya et al., 2001; Nishi et al., 2003; Fig. 8). Type section. The Nakafutamata River section, in the Depositional environment. The mudstone in the northern Tomamae area (Figs. 5, 6; section 7). northern Tomamae area contains abundant benthic for- aminifers, e.g., Gavelinella, Gyroidinoides, Hoeglundina, Stratigraphic relationships. The formation is exposed Oolina, and Silicosigmoilina. The palaeodepth of this in the Tomamae and Mikasa areas, and conformably assemblage is considered to be outer shelf (Sliter and overlies the Saku and Mikasa formations, respectively. Baker, 1972). The cross-stratified sandstone situated It is the synchronous shallower-water facies of the towards the top of each succession (i.e., KY-6) in the Kashima Formation. northern Tomamae area is inferred to have been de- posited in inner shelf to lower shoreface environments Lithology. The lithology and sedimentary cycles of (Toshimitsu, 1985; Wani, 2003). this formation differ between the Tomamae and Mikasa areas. In the Mikasa area the formation forms a single, 3.9.Hakobuchi Formation (new) coarsening-upwards sequence, from bioturbated sandy mudstone to very fine-grained sandstone. Two felsic Definition. This formation is defined by the pre- volcaniclastic turbidite units of the KY-5 and KY-6 dominance of HCS and TCS sandstones and conglom- (Tsukimi Sandstone Member) are intercalated in the erates, and corresponds to the Hakobuchi Group of lower and middle parts, respectively. The KY-5 unit Matsumoto (1942). consists of thick beds of volcaniclastic sandstones with abundant Inoceramus uwajimensis. In the Tsukimi-sawa Type section. Downstream of the Shuparo River Valley section, the KY-5 becomes channel-fill conglom- (dam site of Lake Shuparo) in the Oyubari area (Figs. 4, erates, including well-rounded pebbles of chert, sand- 6; section 15). stone and mudstones and rhyolitic rocks, with abundant fragments of I. uwajimensis. They are high-density tur- Stratigraphic relationships. This formation conform- bidite beds, exhibiting the Si-2 divisions of Lowe (1982). ably overlies the Kashima Formation in the Oyubari The KY-6 unit (Tsukimi Sandstone Member) is 20 m area and the Haborogawa Formation in the north- thick and exposed in the eastern Mikasa area. This unit eastern Mikasa and Tomamae areas. However, the consists mostly of volcaniclastic sandstones with very unconformable relationship between the Haborogawa thin interbedding of dark grey mudstones. The sand- and Hakobuchi formations is recognized in the north- stone beds range from 0.2-2 m thick, and display the western limb of the Sorachi Anticline, north-western S1-Tb-e divisions of Lowe (1982) (Fig. 7N). Mikasa area. The Hakobuchi Formation is overlain by In the Tomamae area, this formation consists of two a disconformity, or an angular unconformity, and is coarsening-upwards sequences. Both begin with strongly covered by deposits younger than the Paleocene, such as bioturbated mudstone, grade into bioturbated muddy the middle-upper Eocene Ishikari Group, containing sandstone, and end in medium- to coarse-grained, cross- coal measures, or the upper Eocene-lower Oligocene laminated sandstone (Fig. 7O). The lower coarsening- offshore-marine Poronai Group. upward sequence incorporates the volcaniclastic marker unit of the KY-5 in the lower part and the KY-6 at the Lithology. As the deposits of the Hakobuchi For- top, respectively. mation form complicated stacking patterns of the 384 R.Takashima et al./Cretaceous Research 25（2004）365-390 third- and fourth-order depositional sequences (DSs), hemipelagic mudstone-dominant units (Fig. 2). The the lithology varies depending on areas and horizons, as Aptian-Albian sequences are laterally similar in all for the Mikasa Formation (Ando, 2003). The number of study areas, comprising the laminated mudstone Soa- sequences differs between sections, but the maximum shibetsugawa Formation, the turbiditic Shuparogawa reaches over 15, including third- and fourth-order DSs. and Maruyama formations, and the mudstone Hikage- They mainly consist of coarsening-upwards facies suc- nosawa Formation, in ascending order. Sedimentary cessions (CUS), a few tens to 100 m thick, of bioturbated structures, lithological associations and sedimentary se- sandy mudstone to HCS/TCS sandstone (Fig. 7P). Often, there is an associated thin, fining-upwards marine Scarce occurrences of macrofossils, and common occur- succession (FUS), not more than a few metres thick, rences of deep-sea trace fossils, suggest abyssal environ- below the CUS. Fluvial conglomerate, sandstone, ments in the Aptian. Benthic foraminifers indicate the mudstone and sometimes coaly beds are intercalated in upper bathyal zone, about 300-600 m in depth, in the the basal part of the DSs. Thick marine conglomerates Albian. and fuvial-channel conglomerates are subordinately de- The Cenomanian-lower Campanian rocks display veloped at several horizons. Felsic tuffs are interbedded lateral variations in lithology, tending to north-westward at a few horizons, and thicken in the northern Mikasa coarsening and shallowing sequences (Fig. 2). The and northern Tomamae areas, where the beds attain 30 uppermost Albian-Turonian Mikasa Formation is and 80 m in thickness, respectively. Compared with the typically composed of shallow-marine, including fluvial, Mikasa Formation, the Hakobuchi Formation is estuarine and outer-shelf, sediments, suggesting rapid characterized by a smaller amount of offshore mudstone uplift in the western margin of the Yezo Forearc Basin. and a larger amount of conglomerate and tuff. The overlying Haborogawa Formation represents deeper facies than the Mikasa Formation in terms of sedimen- Thickness and distribution. The thickness is variable, tary environment, as indicated by the occurrence of ranging from several tens to 450 m thick in the Mikasa shoreface to outer shelf, and partly upper bathyal basin- and Oyubari areas (Ando, 1997). The formation is 250 m plain deposits. On the other hand, the Cenomanian- thick in the northernmost part of the Tomamae area. Turonian Saku Formation and Coniacian-Campanian Kashima Formation exhibit slope fan on continental Fossils and age. Although a few macrofossils, such as slope and basin-plain successions, respectively. Based on Sphenoceramus schmidti (Michael), S. hetonaianus (Mat- benthic foraminiferal assemblages, these formations sumoto), and Inoceramus shikotanensis Nagao and were deposited in upper bathyal depths. Matsumoto, occur in this formation (Ando, 1997; Ando The neritic to non-marine Hakobuchi Formation, et al., 2001), there are no age-diagnostic species. The deposited during the Campanian-Early Maastrichtian planktonic foraminifers Globotruncana rugosa (Marie) and Late Paleocene interval, completely covers the deep- and Subbotina triloculinoides (Plummer) occur in the sea sediments across the whole area from north to south. lower and upper parts of this formation, respectively This formation indicates the final stage of deposition (Yasuda, 1986). The former is assigned to the Campa- and uplift of the Yezo Forearc Basin. nian (Robaszynski et al., 1984), and the latter is of Paleocene age (Yasuda, 1986). The Cretaceous/Palae- ogene boundary sequence has not been detected in the 5. OAE horizons in the Yezo Group Hakobuchi Formation. Laminated, organic-rich marine sediments, indicating Depositional environment. Depending upon the strati- six oceanic anoxic events (OAEs 1a-1d, 2 and the mid- graphic position within the third- or fourth-order DSs, Cenomanian Event, MCE) have been reported from the depositional environments represented change regu- mid-Cretaceous deep-sea cores and land sections in the larly and repetitively. They include shallow-marine envi- European Basin, North and South Atlantic oceans, and ronments, such as outer to inner shelf and shoreface, equatorial mid-Pacific mountains and plateaus (e.g.. and subordinately estuarine, incised valley, and riverine Erbacher et al., 1996; Leckie et al., 2002; Coccioni and gravelly/sandy river-channel, back-marsh and flood- Galeotti, 2003). These sediments are visually-distinctive, plain. so-called “black shales", found as black to dark grey intercalations in white, pelagic limestones or grey marl- stones. These black shale intervals commonly consist 4. Lateral change of depositional environments mainly of finely alternating layers of calcareous fossil- in the Yezo Basin rich laminae and clay/organic matter (Arthur and Sageman, 1994). The Yezo Group is lithologically characterized by However, terrigenous, siliciclastic sequences exposed an alternating sequence of turbidite-dominant and in the North Pacific regions, such as the Yezo Group in R.Takashima et al./Cretaceous Research 25（2004）365-390 385 Japan, are generally composed of dark grey mudstones correlates with the OAE1a event (Fig. 8). The OAEla and pale grey sandstones. Dark grey, monotonously- interval in the study area is very thick, about 110 m coloured deposits throughout the successions prevent (Fig. 6; section 11), compared with 1-5 m in the Italian visual recognition of OAE levels in background sedi- and Swiss sections (Menegatti et al., 1998). The extra- ments in certain outcrops. In the Yezo Group, the ordinarily thick interval in the Yezo Group resulted average total organic carbon (TOC) content is about from a huge terrestrial influx from the active Asian Rb (ppm) continental margin. Mudstone in this interval is charac- those in the Tethyan and Atlantic carbonate rocks and terized by rare bioturbation and fine lamination, with Sr (ppm) scarce or no benthic macro- and microfaunas. organic matter is mostly derived from terrestrial plants throughout the sequences, suggesting that organic frag- 5.2.OAE1b horizon ments (origin of the TOC) were transported from the westward Asian continental margin. Therefore, as the OAElb was a long-lived event, extending from the TOC of the Yezo Group is not a useful indicator of Ticinella bejaouaensis to Hedbergella planispira pf zones anoxic environments, the biostratigraphy and chemo- (114.5-108.21 Ma, about 6.3 myr; Hardenbol et al., stratigraphy are both very important in identifying OAE 1998). This event is represented by several black shales: horizons in this sequence. Several chemostratigraphic (1) the uppermost Aptian Niveau Jacob in the Vocon- studies, using organic carbon isotopes from the Saku &WPT Formation, have reported a positive 813Corg excursion lower Albian Niveau Kilian, Paquier and Leenhardt in record related to the global burial event of marine the Vocontian Basin, and Livellos Monte Nerone and organic matter around the Cenomanian/Turonian boun- Urbino in the Apennines of central Italy. dary (OAE2; Hasegawa and Saito, 1993; Hasegawa, In the Yezo Group, an interval spanning the T. 1995, 1997; Hasegawa and Hatsugai, 2000). Moreover, bejaouaensis to H. planispira pf zones was not detected, another positive 813Corg excursion (ca. 3.6%) was enrichr observed in the Soashibetsugawa Formation of the bed (Fig. 8). Consequently, OAE1b is missing owing to northern part of the Oyubari area (OAEla; Ando et al., erosion or hiatus by a tectonic event in the study area. 2002, 2003). Here we summarize the published bio- stratigraphic and chemostratigraphic results of the study 5.3. OAElc horizon areas using the new stratigraphic classification presented in this paper (Fig. 8), and identify and describe the OAE Black shales representing the OAElc event are horizons in the sequence. characterized by abundant terrigenous organic matter in the lower Upper Albian Biticinella breggiensis pf Zone 5.1. OAEla horizon in central Italy (Amadeus Level) (Erbacher et al., 1996). The 813Ccarbonate excursions within the B. breggiensis pf The late Early Aptian OAEla, spanning about 1 myr Zone are not distinct in the Italian section (Erbacher (119.5-120.5 Ma; Larson and Erba, 1999), named as et al., 1996), whereas a prominent negative excursion of 813Corg (1.5%) has been identified in the Santa Rosa the Niveau Goguel in the Vocontian Basin of France, Livello Selli in Italy and Fischschiefer in Germany, is section, Mexico (Bralower et al., 1999). a global, organic-carbon burial-event. This OAE oc- In the Yezo Group, the B. breggiensis pf Zone spans curred within the basal part of the Leupoldina cabri pf the uppermost part of the Shuparogawa Formation to Zone (Premoli-Silva et al., 1999). There was a microfossil the basal part of the Hikagenosawa Formation (Fig. 8), and is 300 m thick. A negative 813Corg excursion of about E-MORB decrease, and nannoconid (calcareous nannoplankton) 1% occurs at the top of the Shuparogawa Formation species disappeared around the OAEla event (Erba, (Hirano and Fukuju, 1997), and is followed by a radio- 1994; Premoli-Silva et al., 1999). The sharp negative larian high-productivity event in the Maruyama Forma- tion (Fig. 8). This excursion interval could be correlatable abrupt, prolonged positive one (>2%o), are recorded in with the OAElc horizon of Leckie et al. (2002). this event (Menegatti et al., 1998). In the Yezo Group, a pair of negative and positive 5.4. OAEld horizon ss u pn uq sa suxs ro gawa Formation of the Sorachi Group in the Oyubari The uppermost Albian black shales (OAE1d, Breis- -MORB troffer Level) occur in the upper part of the Rotalipora redefined as the Soashibetsugawa Formation of the appenninica pf Zone, as well as in the Stoliczkaia dispar Yezo Group, as described above. Because the paired ammonite Zone, in Tethyan-Atlantic regions (Gale excursion spikes occur near the first appearance datum et al., 1996; Wilson and Norris, 2001). Broad positive (FAD) of Leupoldina cabri, this positive excursion 386 R.Takashimaet al./CretaceousResearch25（2004）365-390 globally significant organic-carbon burial (Erbacher 1982; Arthur and Dean, 1986; Summerhayes, 1987). In et al., 1996; Gale et al., 1996; Nederbragt et al., 2001; central Italy, this event is associated with upper Wilson and Norris, 2001). The OAE1d horizon in the Cenomanian organic-rich layers up to the Cenoma- equatorial Atlantic area is marked by a collapse of nian/Turonian boundary. The MCE level shows a global upper water-column stratification due to intensified win- positive excursion in carbon isotope records, correlatable ter mixing and reduced summer stratification (Wilson within the middle part of the Dicarinella algeriana pf and Norris, 2001). Biological records indicate that this Subzone of the Rotalipora cushmani pf Zone. Above this OAE event damaged planktonic foraminiferal and event, radiolarian and benthic foraminifers decrease in radiolarian populations, as well as the carbonate plat- diversity in central Italy (Erbacher et al., 1996; Damesté forms (Erbacher et al., 1996; Nederbragt et al., 2001). and Koster, 1998; Coccioni and Galeotti, 2003). In the Yezo Group, the R. appeninnica pf Zone is Hasegawa (1997) proved a sudden positive excur- correlatable with the lower middle part of the Hikage- sion of 813Corg (about 1.2%) within KS19a (= the nosawa Formation (Fig. 8). Although the base of the D. algeriana pf Subzone) in the Hakkin-zawa River R. appeninnica pf Zone is indefinite because of the pre- (Shirokin River in Hasegawa, 1997) section of the sence of a thick interval of rare to barren planktonic Oyubari area. This positive peak is probably correlatable foraminifers, mudstones in the lower part of this forma- with the MCE horizon in central Italy. Therefore, the tion contain the upper Albian ammonoids Mortoniceras horizon is placed at the top of the Hikagenosawa For- rostratum and Mariella bergeri (Kawabe et al., 2003; mation of our lithostratigraphic division. Benthic fora- Fig. 8). These ammonoids are index species of the upper miniferal assemblages show a stepwise extinction, marked Albian Stoliczkaia dispar ammonite Zone, as well as the by a 43% reduction of calcareous species during the 3 myr R. appeninnica pf Zone, in Tethyan biozonations. A after the MCE event in the southern Oyubari area (Kaiho broad, positive excursion of the 8l3Corg, which is similar and Hasegawa, 1994; Fig. 8). Radiolarian abundances to the 813Ccarbonate curve of OAEld in the Tethyan- decrease during the same interval (Taketani, 1982). Atlantic region, is observed in the lower part of the Hikagenosawa Formation, just below the Kanajiri 5.6. OAE2 horizon Sandstone Member (KY-3) of the Yezo Group (Hirano and Fukuju, 1997; Fig. 8). This horizon is located within OAE2 (the Bonarelli Event) was the largest environ- the upper part of the Stoliczkaia dispar ammonite mental and biotic disturbance during the mid-Creta- Chronozone (Kawabe et al., 2003). Hence, we concluded ceous interval. Major phenomena include maximum that this excursion level is correlatable with OAE1d. mixed-layer temperatures (about 33-34 °C; Wilson The rocks correlated to the OAEld interval are et al., 2002), an abrupt drop in 87sr/86sr isotopic values weakly-laminated and pyrite-rich mudstones. Planktonic (Bralower et al., 1997), increased volcanic deposits and foraminifers are very rare to absent, whereas spumellar- intensive hydrothermal activity (Sinton and Duncan, ian radiolarians suddenly increased in abundance. 1997). This event is marked by a positive excursion of Agglutinated benthic foraminifers are common, with assemblages of Bathysiphon, Glomospira and Haplo- tacea pf Zone, close to the Cenomanian/Turonian (C/T) phragmoides, and rare Gyroidinoides and Lenticulina. boundary, and is considered to be related to increased The dominance of agglutinated forms might reflect productivity and global expansion of the oxygen relatively dysoxic conditions, as well as the Cenoma- minimum zone (e.g., Damesté and Koster, 1998; nian/Turonian boundary (Kaiho and Hasegawa, 1994). Premoli-Silva et al., 1999). A significant turnover of Although a dissolution effect on calcareous species can- radiolarians, calcareous nannofossils, and deep-dwelling not be ruled out, our results suggest that dysoxic con- planktonic foraminifers occurred, with a high extinction ditions developed and a radiolarian high-productivity rate (26% of all genera) of benthic foraminifers event took place during the OAE1d in the north-west (Erbacher et al.,1996; Erbacher and Thurow,1997; Pacific. Premoli-Silva et al., 1999). The C/T boundary, based on macro- and microfossil 5.5.The mid-Cenomanian Event horizon biostratigraphy, is placed at the base of the Hakkin Muddy Sandstone Member of the Saku Formation in The mid-Cenomanian Event (MCE) is a major the Oyubari and Tomamae areas (Figs. 4, 6; sections 10, 13, and 14). A positive 813Corg excursion of about 2.5% turnover of foraminifers and radiolarians associated is recorded at this boundary, indicating OAE2 (e.g.. shift in the Sr/Ca ratio (Coccioni and Galeotti, 2003). Hasegawa and Saito, 1993). Investigation of benthic The organic-rich layers corresponding to this event have foraminiferal assemblages in the Oyubari area demon- been found in central Italy, the Vocontian Basin strates that the two lowest oxygen conditions devel- (Coccioni and Galeotti, 2003), and the North Atlantic oped during 0.15 myr before the C/T boundary and (Deep Sea Drilling Project Sites 367, 386 and 398; Cool, 1.5 myr after (Kaiho and Hasegawa, 1994). Radiolarians R.Takashima et al./Cretaceous Research 25（2004）365-390 387 become abundant just above the C/T boundary (e.g., for support during our fieldwork. We thank T. Hatsugai, Hasegawa, 1997; Fig. 8). A. Ennyu, S. Egawa, F. Suzuki, A. Ando and T. Sakai for provision of unpublished data. This study was supported financially by the JSPS Research Fellows 6. Conclusions Scheme (No. 09898 to Takashima, No. 06365 to Wani, and No. 07622 to Moriya), Grants-in-Aid from JSPS Our study of the Yezo Group, a thick, forearc basin- (No. 14740302 to Kawabe, No. 10640446 to Ando, fill sequence, distributed across Hokkaido, Japan, serves and Nos. 13354006 and 15340176 to Nishi), and in part as an important record of tectonic, faunal and en- by a 21st Century Center of Excellence (COE) Program vironmental change on the north-west Pacific margin on “Neo-Science of Natural History'” at Hokkaido during the Cretaceous Period. The litho- and bio- University. stratigraphic summaries of the sequence were previously little known, and various lithostratigraphic divisions had been proposed. Based on our wide areal mapping of the sequences, as well as analyses of litho- and biofacies, References we have here proposed a new lithostratigraphic scheme for the sequence. Ando, H., 1990a. Shallow-marine sedimentary facies distribution and The Aptian-Albian sequences (the Soashibetsugawa, progradational sequences of the Mikasa Formation, Middle Yezo Shuparogawa, Maruyama and lower Hikagenosawa Group (Upper Cretaceous). Journal of the Geological Society of formations) of the group were deposited on a continental Japan 96, 453-469 (in Japanese, English abstract). slope with a lower-upper bathyal palaeodepth in all Ando, H., 1990b. Stratigraphy and shallow marine sedimentary facies sections. The Cenomanian to lower Campanian sequen- of the Mikasa Formation, Middle Yezo Group (Upper Creta- ceous). Journal of the Geological Society of Japan 96, 279-295 ces exhibit lateral variations of sedimentary environ- (in Japanese, English abstract). ment in the study areas. Generally, the sequences Ando, H., 1997. Apparent stacking patterns of depositional sequences exposed in the Mikasa and Tomamae areas show in the Upper Cretaceous shallow-marine to fuvial successions, shallower facies of shelf to non-marine environments northeast Japan. Memoir of the Geological Society of Japan 48, 43-59. (the Mikasa and Haborogawa formations), while those Ando, H., 2003. Stratigraphic correlation of Upper Cretaceous to in the Oyubari area represent continental slope environ- Paleocene forearc basin sediments in northeast Japan: cyclic ments (the upper Hikagenosawa, Saku and Kashima sedimentation and basin evolution. Journal of Asian Earth formations). Above the lower Campanian, the neritic to Sciences 21, 921-935. non-marine facies of the Hakobuchi Formation were Ando, H., Tomosugi, T., Kanakubo, T., 2001. Upper Cretaceous to Paleocene Hakobuchi Group, Nakatonbetsu area, northern deposited in all areas. Hokkaido—lithostratigraphy and megafossil biostratigraphy. s s Journal of the Geological Society of Japan 107, 142-162. shales, as seen in the Tethyan and Atlantic regions, Ando, A.,Kakegawa, T.,Takashima, R.,Saito, T., 2002. .New because of rapid sedimentation and a constant volumi- perspective on Aptian carbon isotope stratigraphy: data from 813C records of terrestrial organic matter. Geology 30, 227-230. nous influx of terrigenous detritus. However, the inte- Ando, A., Kakegawa, T., Takashima, R., Saito, T., 2003. Stratigraphic gration of the bio- and chemostratigraphic results of the carbon isotope fuctuations of detrital woody materials during the study areas, using our new stratigraphic scheme, reveals Aptian Stage in Hokkaido, Japan: comprehensive 8l3C data from the OAEla, 1c, 1d, MCE and OAE2 horizons in the four sections of the Ashibetsu area. Journal of Asian Earth Yezo Group. The OAEla, 1c, 1d and 2 horizons are Sciences 21, 835-847. Arthur, M.A., Dean, W.E., 1986. 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American Geophysical Union, pp. 5-37. We express sincere gratitude to Y. Takayanagi, T. Bralower, T.J., Arthur, M.A., Leckie, R.M., Sliter, W.V., Allard, D.J., Saito (Tohoku University), H. Hirano (Waseda Uni- Schlanger, S.O., 1994. Timing and paleoceanography of oceanic versity), S. Sano (Fukui Prefectural Dinosaur Museum), dysoxia/anoxia in the Late Barremian to early Aptian (Early T. Tomosugi (Kyushu University), and K. Kurihara Cretaceous). Palaios 9, 335-369. Bralower, T.J., Fullagar, P.D., Paull, C.K., Dwyer, G.S., Leckie, (Mikasa City Museum) for helpful suggestions, and R.M., 1997. Mid-Cretaceous strontium-isotope stratigraphy of also to K. Kobayashi, K. Hasegawa, S. Matsuda, deep-sea sections. Geological Society of America, Bulletin 109, K. Matsuda, N. Watanabe, K. Komoto (Mikasa City) 1421—1442. 388 R.Takashima et al./Cretaceous Research 25（2004）365-390 Bralower, T.J., Cobabe, E., Clement, B., Sliter, W.V., Osburn, C.L., Hirano, H., Fukuju, T., 1997. 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Cretaceous Research 17, 515-606. Japan. Monograph of the Association of Geological Collaborators Girard, M., Jolivet, L., Nakagawa, M., Aguirre, L., Nida, K., 1991. of Japan 31, 1-15 (in Japanese, English abstract). Acidic volcanic products in Lower Cretaceous deposits of the Kito, N., Kiminami, K., Nida, K., Kanie, Y., Watanabe, T., Sorachi-Yezo Belt, Hokkaido, northeast Japan. Journal of the Kawaguchi, M., 1986. The Sorachi Group and the Yezo Super- Geological Society of Japan 97, 1-14. group: late Mesozoic ophiolites and forearc sediments in the axial Hardenbol, J., Thierry, J., Farley, M.B., Jaquin, T., de Graciansky, zone of Hokkaido. Monograph of the Association of Geological P.-C., Vail, P.R., 1998. Cretaceous chronostratigraphy. In: Collaborators of Japan 31, 81-96 (in Japanese, English abstract). Graciansky, P.-C., Hardenbol, J., Jaquin, T., Vail, P.R. (Eds.), Kodama, K., Maeda, H., Shigeta, Y., Kase, T., Takeuchi, T., 2002. Mesozoic and Cenozoic Sequence Chronostratigraphic Framework Integrated biostratigraphy and magnetostratigraphy of the Upper of European Basins. SEPM (Society for Sedimentary Geology) Cretaceous System along the River Naiba in southern Sakhalin, Special Publication 60, 3-13. Russia. Journal of the Geological Society of Japan 108, 366-384 Hasegawa, T., 1995. Correlation of the Cenomanian/Turonian (in Japanese, English abstract). boundary between Japan and Western Interior of the United Larson, R.L., 199la. Geological consequences of superplumes. States. Journal of the Geological Society of Japan 101, 2-12. Geology 19, 963-966. Hasegawa, T., 1997. Cenomanian-Turonian carbon isotope events Larson, R.L., 1991b. Latest pulse of Earth: evidence for a mid-Creta- recorded in terrestrial organic matter from northern Japan. ceous superplume. Geology 19, 547-550. Palaeogeography, Palaeoclimatology, Palaeoecology 130, 251-273. Larson, R.L., Erba, E., 1999. Onset of the mid-Cretaceous greenhouse Hasegawa, T., Hatsugai, T., 2000. Carbon-isotope stratigraphy and its in the Barremian-Aptian: igneous events and the biological, chronostratigraphic significance for the Cretaceous Yezo Group, sedimentary, and geochemical responses. Paleoceanography 14, Kotanbetsu area, Hokkaido, Japan. Paleontological Research 4, 663-678. 95-106. Leckie, R.M., Bralower, T.J., Cashman, R., 2002. Oceanic anoxic Hasegawa, T., Saito, T., 1993. Global synchroneity of a positive events and plankton evolution: biotic response to tectonic forcing carbon isotope excursion at the Cenomanian/Turonian boundary: during the mid-Cretaceous. Paleoceanography 17, 129. validation by calcareous microfossil biostratigraphy of the Yezo Lowe, D.R., 1982. Sediment gravity flows: II. Depositional models Group, Hokkaido, Japan. The Island Arc 3, 181-191. with special reference to the deposits of high-density turbidity Hashimoto, W., 1953. Explanatory text of the geological map of Japan currents. Journal of Sedimentary Petrology 52, 279-297. “Yamabe' (scale 1:50000). Geological Survey of Hokaido 82 pp. Maeda, H., 1986. Stratigraphy and fossil molluscan fauna of the (in Japanese, English summary). Upper Cretaceous System in the Bibai area, Hokkaido, north Hirano, H., 1995. Correlation of the Cenomanian/Turonian boundary Japan. Journal of the Geological Society of Japan 92, 135-153 (in between Japan and Western Interior of the United States in Japanese, English abstract). relation with oceanic anoxic events. Journal of the Geological Matsumaru, K., 1971. Certain larger foraminifera from Japan. Journal Society of Japan 101, 13-18. of Saitama University, Faculty of Education 20, 149-159. R.Takashima et al./Cretaceous Research 25（2004）365-390 389 Matsumoto, T., 1942. Fundamentals in the Cretaceous stratigraphy of Sinton, C.W., Duncan, R.A., 1997. Potential links between ocean Japan, Part 1. Memoirs of the Faculty of Science, Kyushu plateau volcanism and global ocean anoxia at the Cenomanian/ University, Series D, Geology 1, 129-280. Turonian boundary. Economic Geology 92, 836-842. Matsumoto, T., 1951. The Yezo Group and the Kwanmon Group. Sliter, W.V., 1989. Aptian anoxia in the Pacific Basin. Geology 17, Journal of the Geological Society of Japan 57, 95-98 (in Japanese, 909-912. English abstract). Sliter, W.V., Baker, R.A., 1972. Cretaceous bathymetric distribution Matsumoto, T., 1977. Zonal correlation of the Upper Cretaceous in of benthic foraminifers. Journal of Foraminiferal Research 2, Japan. Palaeontological Society of Japan, Special Paper 21, 63-74. 167-183. Matsumoto, T., Okada, H.,1973. Saku Formation of the Yezo Summerhayes, C.P., 1987. Organic-rich Cretaceous sediments from the geosyncline. Science Report Department of Geology, Kyushu North Atlantic. In: Brooks, J., Fleet, A.J. (Eds.), Marine Petroleum University 11, 275-309. Source Rocks. Geological Society, London, Special Publication 26, Menegatti, A.P., Weissert, H., Brown, R.S., Tyson, R.V., Farrimond, 301-316. P., Straser, A., Caron, M., 1998. High-resolution 83C stratigra- Takashima, R., Nishi, H., Saito, T., Hasegawa, T., 1997. Geology and phy through the early Aptian *"Lievello Selli?' of the Alpine Tethys. planktonic foraminiferal biostratigraphy of Cretaceous strata Paleoceanography 13, 530-545. distributed along the Shuparo River, Hokkaido, Japan. Journal Moriya, K., Hirano, H., 2001. Cretaceous stratigraphy inthe of the Geological Society of Japan 103, 543-563 (in Japanese, Chikubetsu area, Hokkaido. Journal of the Geological Society of English abstract). Japan 107, 199-214 (in Japanese, English abstract). Takashima, R., Yoshida, T., Nishi, H., 2001. Stratigraphy and Moriya, K., Nishi, H., Tanabe, K., 2001. 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Taphofacies models for Upper Cretaceous ammo- Leupoldina cabri (Sigal, 1952) from the Sorachi Group of noids from the Kotanbetsu area, northwestern Hokkaido, Hokkaido, Japan and its bearing on the Cretaceous chronology Japan. Palaeogeography, Palaeoclimatology, Palaeoecology 199, of northwestern Pacific marine strata. Bulletin of the National 71-82. Science Museum, Series C, Geology 26, 183-191. Wani, R., Hirano, H., 2000. Upper Cretaceous biostratigraphy in the Sano, S., 1995. Litho- and biofacies of Early Cretaceous rudist-bearing Kotanbetsu area, northwestern Hokkaido. Journal of the Geo- carbonate sediments in northeastern Japan. Sedimentary Geology logical Society of Japan 106, 171-188 (in Japanese, English 99, 179-189. abstract). 390 R.Takashima et al./Cretaceous Research 25（2004）365-390 Wilson, P.A., Norris, R.D., 2001. Warm tropical ocean surface and from the core of the Turonian tropics on Demerara Rise. Geology global anoxia during the mid-Cretaceous period. 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J. Phys. Earth, 41, 165-188, 1993
Seismic Refraction Study in the Kitakami Region,
Northern Honshu, Japan
Takaya Iwasaki,1,õ Toshikatsu Yoshii,1 Takeo Moriya,2
Akio Kobayashi,3 Makoto Nishiwaki,4 Tomoki Tsutsui,5,*
Takashi Iidaka,1 Akira Ikami,6,**
and Tetsu Masuda7,***
1Earthquake Research Institute, the University of Tokyo, Bunkyo-ku,
Tokyo 113, Japan
2Department of Geophysics, Faculty of Science, Hokkaido University,
Sapporo 060, Japan
3 Matsushiro Seismological Observatory, Japan Meteorological Agency,
Nagano 381-12, Japan
4 Japan Meteorological Agency, Chiyoda-ku, Tokyo 100, Japan
5Geophysical Institute, Faculty of Science, Kyoto University,
Kyoto 606, Japan
6Research Center for Seismology and Volcanology, School of Science,
Nagoya University, Nagoya 464-01, Japan
7 Geophysical Institute, Faculty of Science, Tohoku University,
Sendai 980, Japan
An extensive seismic refraction experiment with the use of explosive sources was
conducted in the Kitakami region, northern Honshu, Japan, on November 1, 1990. The
experiment area is divided into two geological units by the Hayachine Tectonic Belt
(HTB). The southern terrane consists of pre-Silurian basements and Silurian-lower
Cretaceous marine sediments, while the northern one is characterized by a Jurassic
accretionary complex. Both of the units were intruded by Cretaceous granitic rocks.
An almost N-S seismic refraction profile of 194-km length was extended from Kuji
City, Iwate Prefecture to Ishinomaki City, Miyagi Prefecture, on which 4 shots with a
charge size of 450-700 kg were fired. The generated seismic signals were recorded at
179 stations, from which a precise crustal structure was determined. The uppermost
crust is covered with a very thin (0.5-1 km) surface layer with a velocity of 3.1-5.4 km/s.
Our results show a remarkable structural difference between the northern and southern
Kitakami terranes. A velocity of the uppermost crust is 5.90 km/s in the northern part
of the profile while 6.05-6.15 km/s in the southern part. This lateral velocity change
Received April 5, 1993; Accepted July 12, 1993
õ To whom correspondence should be addressed.
Present address: * Aso Volcanological Laboratory, Faculty of Science, Kyoto University,
Kumamoto 869-14, Japan. ** Earthquake Research Institute, the University of Tokyo,
Bunkyo-ku, Tokyo 113, Japan (passed away on October 1, 1991). *** TOHOKU BRANCH,
OYO CORPORATION, Sendai 983, Japan.
165
166 T. Iwasaki et al.
occurs just beneath the HTB. A mid-crustal interface corresponding to well-recorded
wide-angle reflections (PiP phase) shows an abrupt southward depth decrease from 25
to 20 km beneath the HTB. The reflections from the Moho boundary (PmP phase) were
well observed in all of the record sections. The Moho depth determined from the PmP
phase also decreases southward from 34 to 32 km. The velocity of the lower crust is
6.8-7.0 km/s. The refracted waves from the uppermost mantle (Pn phase) were observed
with weak amplitudes at offsets greater than 160 km. These data suggest that the Pn
velocity in the experiment area is less than 7.7 km/s.
The VP/VS ratio within the crust was determined from S-waves recognized in the
vertical component seismograms. The VP/VS ratio in the upper crust also shows a
significant lateral change at the HTB, namely 1.72-1.73 in the northern part whereas
1.75-1.76 in the southern part. In the lower crust, the VP/VS ratio is 1.73-1.76. In the
present experiment, weak refracted waves from the upper mantle (Sn phase) were
observed at offsets greater than 170 km, from which the Sn velocity was estimated to
be 4.3 km/s.
1 . Introduction
Since 1979, the Research Group for Explosion Seismology (RGES) has carried out
seismic refraction experiments in the Japanese Islands under the Japanese Earthquake
Prediction Program. Most of the experiments in a period of 1979-1988 focused on the
investigation on the uppermost crustal structures from rather short profiles of 50-70 km
(Asano et al., 1979; Yoshii et al., 1985, 1986; Sasatani et al., 1990; Matsu'ura et al.,
1991; RGES, 1992a, b). Seismic refraction studies in the 6-th stage of the program,
which started in 1989, are aimed to reveal the detailed structures of lower crust and
upper mantle by shooting relatively long (150-200 km) profiles. The second experiment
of this stage was conducted in the Kitakami region, northern Honshu, Japan, in the
autumn of 1990 (Fig. 1).
Northern Honshu is one of the typical island arcs located in the western rim of
the Pacific Ocean. The Pacific Plate starts to be subducted westward beneath the Japan
Trench about 200 km east of northern Japan. The plate subduction under northern
Honshu is well traced by the double seismic planes determined from a dense seismic
network (Hasegawa et al., 1978). The Kitakami region forms the eastern part of northern
Honshu, whose western and eastern sides are bordered by the volcanic front and the
aseismic front (Yoshii, 1975), respectively. The Kitakami region is geologically divided
into two terranes (Editorial Committee of Tohoku, 1989). The southern terrane is
composed of pre-Silurian basements and Silurian-lower Cretaceous sedimentary rocks.
The northern one, on the other hand, is an accretionary complex of Jurassic sediments.
Both of the units experienced extensive intrusions at the Cretaceous time. The boundary
region between these units is called the Hayachine Tectonic Belt (HTB) characterized
by mafic to ultramafic rocks of the Hayachine Complex.
The velocity structure in northern Honshu was investigated by several authors.
Matsuzawa (1959) determined the velocity structure from a series of seismic refraction
studies (Ishibuchi-Kamaishi Profile, see Fig. 1). His crustal model consists of three layers
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 167
Fig. 1. Geological division of the Kitakami region, northern Honshu, Japan,
modified after Geological Survey of Japan (1992). Major tectonic boundaries
around Japan are presented in the index map, where our studied area is
indicated by broken lines. Lines "A" and "B" denote Ishibuchi-Kamaishi
Profile (Matsuzawa, 1959) and Oga-Kesennuma Profile (Hashizume et al., 1968;
Yoshii and Asano, 1972), respectively. The Kitakami region is divided into
two geological blocks (northern and southern Kitakami terranes) by the
Hayachine Tectonic Belt (HTB) running in NW-SE direction. Both blocks were
intruded by the Cretaceous granitic rocks. The HTB is characterized by mafic
to ultramafic rocks. a, Kuzumaki Fault; b, Eastern Boundary Fault of the
Hayachine Tectonic Belt; c, Hizume-Kesennuma Fault.
Vol. 41, No. 3, 1993
168 T. Iwasaki et al.
of 2.5, 5.8, and 6.1-6.2 km/s, respectively. The velocity of the uppermost mantle (Pn
velocity) obtained is 7.5-8.0 km/s. The upper boundary of the 6.1-6.2 km/s layer becomes
shallower to the east, and its depth is only 0.5-2.5 km beneath the southern Kitakami
terrane. The refraction experiment along Oga-Kesennuma Profile (Fig. 1) revealed for
the first time the detailed velocity structure of northern Honshu (Hashizume et al.,
1968; Yoshii and Asano, 1972). This profile was extended from the continental slope
off the Pacific coast to the back arc basin in the Japan Sea crossing northern Honshu
with NW-SE direction. The velocities of the upper and lower crusts were estimated to
be 5.9 and 6.6 km/s, respectively, although the latter value was poorly resolved. The
crustal thickness is about 30 km in almost the entire part of the arc but it shows a
remarkable thinning under the Pacific Ocean and the Japan Sea. The Pn velocity has
a strong lateral change, namely, 7.5 km/s under the northern Honshu whereas 8.2 km/s
under the Japan Sea. The rather low value of 7.5 km/s was confirmed by the further
seismic refraction experiments with the use of large underwater explosions in the Pacific
Ocean (Okada et al., 1979; Asano et al., 1979; Yoshii et al., 1981).
The velocity structure in the northern Honshu was also obtained from the travel
time inversion for natural earthquakes (Zhao, 1988; Zhao et al., 1990). Minoura and
Hasegawa (1992) modified their structures and presented maps of upper crustal velocity,
and Conrad and Moho discontinuities in the northern Honshu. According to these
maps, the upper crustal velocity increases eastward from 5.9 to 6.1 km/s, while the
velocity in the lower crust is 6.6 km/s almost in the entire region. The Conrad and
Moho boundaries show concave geometries whose deepest parts are located at 20- and
36-km depths in the central part of northern Honshu, respectively. The Pn velocity in
their structure is 7.5 km/s, which is in good agreement with that from the Oga-Kesennuma
Profile.
2. Experiment and Data Processing
The seismic refraction experiment was conducted on November 1, 1990 (RGES,
1992c). A 194-km profile was extended in N-S direction from Kuji City, Iwate Prefecture
to Ishinomaki City, Miyagi Prefecture (Fig. 2). On this profile, 4 explosive shots were
fired, whose locations, charge sizes and shot times are listed in Table 1. It is noted that
the shot point S-2 was located quite close to the HTB (Figs. 1 and 2). Figure 2 also
shows 179 observation stations temporally deployed with an average spacing of 1.2 km.
The locations and observers of these stations were reported by RGES (1992c). Sensors
Table 1. Shot times, locations and charge sizes of four explosions in the 1990
Kuji-Ishinomaki Profile experiment.
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 169
Fig. 2. Location map of 4 shot points (crosses) and 179 observation stations (solid
circles). The seismic stations were deployed with a spacing of 1.2 km on a
194-km profile.
we used were geophones (L-22D, Mark Products Co., Ltd) with a natural frequency
of 2.0-Hz. As for the recording instruments in the experiment, digital recorders were
deployed at 98 stations. At the other 81 stations, compact analogue cassette recorders
(Yoshii, 1980) were used to collect the seismic signals. In the vicinity of each shot point,
6 additional observing points were set with a spacing of 100 m for the direct measure-
ment of surface velocity. In either case of digital or analogue recording system, the total
frequency response was kept flat in a frequency range of 0.5-30 Hz. Clocks at the
observation stations were carefully checked against the Japan Standard Time (JJY),
which ensured a timing accuracy of 5-10 ms in our seismic data.
Vol. 41, No. 3, 1993
170 T. Iwasaki et al.
The data collected by the analogue recorders were AD converted after anti-alias
filtering with a sampling frequency of 100 Hz on a work station, HP350H (Hewlett-
Packard Company), Earthquake Research Institute, the University of Tokyo. These
data were corrected for timing and fluctuation of tape speed in the recorder system
and stored in a standard format. The seismic data by the digital recorders were also
converted to the same standard format after the correction of timing on a personal com-
puter (PC-9801, NEC Corp.), and sent to HP350H for the final editorial work.
3. Characteristics of Observed Seismograms
Figure 3(a) and (b) shows original record sections for S-1 and S-4, respectively. In
general, the quality of the data is quite good and clear first arrivals corresponding to
a refracted wave from the upper crust (Pg phase) can be traced up to offsets of
160-170 km. The intercept times of the Pg phase are only 0.1-0.5 s, indicating the surface
layer is very thin along the profile. In the section of S-1 (Fig. 3(a)), the apparent velocity
of the Pg phase is 5.9-6.0 km/s within an offset of 70-80 km. The Pg phase from S-4,
on the other hand, has a higher velocity of 6.1-6.2 km/s in the same offset range (Fig.
3(b)). As described in the next section, this reflects a lateral structural difference between
the northern and southern Kitakami terranes. Weak arrivals at offsets greater than
160-180 km have an apparent velocity greater than 7.0 km/s. These are interpreted as
a refracted wave from the uppermost mantle (Pn phase).
Many later phases are recognized in all of the sections. In Fig. 3(a), clear phases,
R1 and R2, at offsets of 60-150 km are wide-angle reflections from a mid-crustal interface
(PiP phase) and the Moho boundary (PmP phase), respectively. These phases are
enhanced by the low-pass filter of 15-20 Hz. The other two phases, L1 and L2, which
are seen in a limited offset range of 30-60 km, contain rather high frequency components
as compared with R1 and R2. As described in Sec. 4.3, we interpret these as reflections
from local interfaces within the crust. The record section of S-4 is also characterized
by several later phases (R3-R5 and L3-L4, see in Fig. 3(b)). A phase recognized at
offsets of 10-40 km (R3) is a reflected wave from a shallow crustal interface while a
phase of R4 corresponds to a reflection from a depth almost comparable to the case
of R1 . A PmP phase is seen at offsets greater than 80 km (R5). Phases of L3 and L4
are localized reflections. Similar localized phases are also found in the sections of S-2
and S-3.
Relatively low frequency phases with a velocity of about 3.5 km/s are found in a
short offset range (0-40 km). These are S-waves and considerably enhanced by low-pass
filtering (¬6 Hz). Observations of such S-waves were also reported from seismic
refraction studies in the central part of Japan (Sasatani et al., 1990; Matsu'ura et al.,
1991).
Travel times of the first arrival and prominent later phases were carefully picked
up on the personal computer PC-9801 and graded with respect to their quality. We
assigned three ranks-A, B and C- which indicate reading errors of 0.01, 0.03 and
0.05 s, respectively.
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 171
(a)
(b)
Fig. 3. Examples of record sections of S-1 and S-4. (a) Record section of S-1.
Each trace is normalized by its maximum amplitude. Reduction velocity is
taken to be 6.0 km/s. Clear first arrivals are recognized up to an offset of
160-170 km. Prominent later phases (R1 and R2) appear above an offset of
60 km. Localized later phases (L1 and L2) are shown in an offset range of
30-60 km. See text for explanation. (b) Record section of S-4. The first arrival
shows a high apparent velocity of 6.1-6.2 km/s. Remarkable later phases are
indicated by R3-R5. L3 and L4 are localized later phases. See text for
explanation.
Vol. 41, No. 3, 1993
172 T. Iwasaki et al.
4. Construction of P-Wave Velocity Structure Model
4.1 Upper crustal structure by time-term method
Surface velocities beneath S-1 to 4 were determined as 3.1, 3.9, 4.3 and 5.4 km/s,
respectively, from the additional measurements conducted in the vicinities of the
individual shot point.
The geometry of the surface layer and the basement velocity were determined by
a time-term method. This method, which was developed by several authors (Scheidegger
and Willmore, 1957; Willmore and Bancroft, 1960; Mereu, 1966), is a conventional but
Fig. 4. Uppermost crustal structure determined by time-term analysis. (a)
Topography along the profile. (b) Time-terms. The open circles are time-terms
obtained assuming a profile consists of four segments with a different basement
velocity. For a comparison, time-terms obtained for a case of a uniform
basement velocity are given by solid circles. (c) Geometry of surface layer and
basement velocities. The depth down to the basement was calculated from the
time-terms with a surface layer velocity of 4.2 km/s. The basement velocity
shows a remarkable lateral change beneath S-2 at which the HTB intersects
our profile.
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 173
Fig. 5. Comparison between observed (solid circles) and calculated travel times
(open circles) in the time-term analysis. Horizontal axis shows an offset from
an individual shot point. The model given in Fig. 4 explains the observed first
arrival time data within 0.05-0.1 s.
Vol. 41, No. 3, 1993
174 T. Iwasaki et al.
efficient inversion method for the determination of the uppermost crustal structure, and
has been widely applied to the seismic refraction studies in the Japanese Islands (e.g.,
Yoshii and Asano, 1972; Asano et al., 1982; Ikami et al., 1986). In the present analysis,
the algorithm of the time-term method was modified to detect the lateral change in
basement velocity. Namely, we divided our profile into four segments according to the
(a)
(b)
Fig. 6
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 175
geological map (Fig, 1), and determined their basement velocities as well as time-terms
beneath the observation points on the basis of the inversion theory developed by Jackson
and Matsu'ura (1985) and Yabuki and Matsu'ura (1993). A data set used for the analysis
consisted of the travel time data of the first arrival within an offset of 60-80 km.
The result by this analysis is presented in Fig. 4. Figure 4(a) shows the topography
along the profile. The obtained time-terms are plotted by open circles in Fig. 4(b). In
this panel, we also present the time-terms obtained under the assumption of a uniform
basement velocity (solid circles). The two sets of the time-terms are almost consistent
with each other except in the northern edge of the profile where the solution of the
inversion was unstable due to the poor data coverage. The geometry of the surface
layer is presented in Fig. 4(c) where a constant surface velocity of 4.2 km/s is assumed
along the profile. Our result shows that the surface layer is only 0.5-1.0 km thick.
Particularly, it is very thin (0.1-0.2 km) in distance ranges of 70-85, 110-130 and
185-194 km measured from S-1. The velocities obtained for the four segments show a
remarkable lateral change. The velocity in the south of S-2, namely the HTB, is
0.1-0.2 km/s higher than that in the northernmost profile. The estimation errors of these
velocities were less than 0.01-0.02 km/s, which indicates that our seismic data resolved
well the lateral heterogeneity in the uppermost crustal structure. We obtained a result
of 6.06 km/s for the case of a constant basement velocity, which is regarded as an average
uppermost crustal velocity in the experiment area. In Fig. 5, travel times calculated
from the obtained time-terms and basement velocities are compared with the observed
data. Our model explains most of the observed data within an error of 0.05-0.10 s.
4.2 Velocity model by ray-tracing method
A detailed velocity structure below the surface layer was mainly determined by a
2-D ray-tracing method. Program packages we used were developed by Cerveny and
Psencik (1983) and Iwasaki (1988). In this process, the amplitude infomation was also
taken into account by computing synthetic seismograms.
The shallower part of the crust was well constrained by the travel times of the Pg
phase. The velocity structure down to a 16-km depth is shown as a contour map in
Fig. 6(a). The geometry of the 3.1 km/s layer in the northern end of the profile was
determined from the travel time data near S-1 and the amplitude behaviors of the Pg
phases from S-1 and S-2. The interface at a depth of 4-7 km was necessary to explain
an abrupt change in apparent velocity of the first arrival observed at offsets of 60-80 km
Fig. 6. Velocity structure model by 2-D ray-tracing method. The origin of the
distance axis is taken at S-1. Shot positions are denoted by solid triangles at
the top of each panel. (a) Contour map of the upper crustal velocity. Note a
remarkable velocity change beneath the HTB. See the text for explanation. (b)
Velocity structure of the whole crust. An asterisk indicates an assumed value.
Numerals in parentheses are VP/VS ratios obtained from the analysis of the
S-wave data. Shaded areas denote localized reflectors within the crust, whose
velocities are 0.3-0.4 km/s less than the surrounding materials. The structure
up to the Moho boundary shows a remarkable lateral change at the HTB. See
the text for explanation.
Vol. 41, No. 3, 1993
176 T. Iwasaki et al.
(a)
(b)
(c)
Fig. 7
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 177
(d)
Fig. 7. Comparisons between observed and calculated travel times. The observed
data are indicated by the open circles. The radius of a circle denotes the rank
of data quality, namely, large, middle and small circles mean the reading errors
of 0.01, 0.03 and 0.05 s, respectively. Solid lines indicate travel time curves of
the predominant phases predicted from our model. Ray-paths for these phases
are shown in Fig. 8. (a) Travel time plot of S-1. (b) Travel time plot of S-2.
(c) Travel time plot of S-3. (d) Travel time plot of S-4.
Fig. 8. Ray-diagram at S-1. Rays for the predominant phases are shown.
Vol. 41, No. 3, 1993
178 T. Iwasaki et al.
in the record sections of S-1 and S-4 (Fig. 3). The later phase of R3 in Fig. 3(b) was
consistently interpreted as a wide-angle reflection from this interface, which is hereafter
denoted as a PiP, phase. The map clearly shows a lateral velocity change in the upper
crust. Namely, the velocity beneath the sediment is 5.90 km/s in the north of the HTB,
while 6.05-6.15 km/s further to the south. Such a southward velocity increase is
recognized down to a depth of 15 km, indicating that the HTB is a prominent boundary
between the northern and southern Kitakami terranes. In Fig. 7, travel time curves
computed from our model are shown with the observed data. It is seen that our model
explains the observed first arrival time data within an error of 0.05-0.1 s.
(a)
(b)
Fig. 9. Observed and synthetic seismograms of S-1. Localized later phases have
not been modelled. Our model well explains the gross features of observed
seismograms. Prominent later phases in an offset range of 60-150 km are
interpreted as PiP2 and PmP phases. See Fig. 7(a) and Fig. 8 for phase
identification. (a) Observed seismograms. Each trace plotted has been digitally
band-pass filtered (3-17 Hz) and normalized by its maximum amplitude. (b)
Synthetic seismograms.
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 179
The structure of the deeper part of the crust was obtained from the travel time
and amplitude data of the clear later phases such as R1, R2, R4 and R5 in Fig. 3. The
modelling for the localized later phases (e.g., L1-L4 in Fig. 3) is described in Sec. 4.3.
Figure 6(b) shows our final crustal model. In this model, the velocity in a depth range
of 15-25 km is 6.45-6.65 km/s. This value was not constrained well by our travel time
data, but could not be larger than 6.6-6.7 km/s. If we assume a velocity of 6.7-6.8 km/s,
for example, the first arrivals at offsets greater than 100-120 km become 0.1-0.2 s earlier
than the observation. The later phases of R1 and R4 in Fig. 3 were well interpreted as
a reflection from a mid-crustal interface at a depth of 20-25 km. Hereafter, we denoted
this reflection as a PiP2 phase. The velocity contrast at the interface was determined as
0.15-0.35 km/s from the amplitudes of the PiP2 phase. It is noted that the depth of this
(a)
(b)
Fig. 10. Observed and synthetic seismograms of S-2. A later phase in the
northernmost part of the profile is interpreted as a PiP2 phase. In the southern
part of the profile, PiP2 and PmP phases are also recognized. See Fig. 7(b)
and Fig. 8 for phase identification. (a) Observed seismograms. (b) Synthetic
seismograms.
Vol. 41, No. 3, 1993
180 T. Iwasaki et al.
interface abruptly changes beneath the HTB.
The strong phases of R2 and R5 are reflections from the Moho boundary (PmP
phase). Such a PmP phase is also seen in the sections of S-2 and S-3. A weak phase at
offsets greater than 160-180 km (Fig. 3) was interpreted as a Pn phase. Although the
onset of this phase is not so clear, its apparent velocity is less than 7.7 km/s. In the
present analysis, we fixed the Pn velocity at 7.5 km/s based on the other seismic refraction
experiments in northern Honshu (Yoshii and Asano, 1972), and determined the lower
crustal structure both from the travel times and amplitudes of the PmP phase. The
obtained velocity between the mid-crustal interface and the Moho increases downward
from 6.8-6.9 to 7.0 km/s (Fig. 6(b)). The Moho depth shows an southward decrease
from 34 to 32 km.
(a)
(b)
Fig. 11. Observed and synthetic seismograms of S-3. A later phase appearing in
a wide offset range of -110 to -60 km is interpreted as PiP2 phase. A PmP
phase is also recognized with a large amplitude at offsets of -130 to -90 km.
See Fig. 7(c) and Fig. 8 for phase identification. (a) Observed seismograms.
(b) Synthetic seismograms.
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 181
(a)
(b)
Fig. 12. Observed and synthetic seismograms of S-4. Later phases in offset ranges
of -130 to -50 km and -190 to — 80 km are well explained as PiP2 and
PmP phases, respectively. See Fig. 7(d) and Fig. 8 for phase identification. (a)
Observed seismograms. (b) Synthetic seismograms.
The travel times of later phases predicted from our final model are also compared
with the observed data in Fig. 7. To show ray-paths of predominant phases, the
ray-diagram for S-1 is presented in Fig. 8. It is seen that the most of the later phases
are explained within an error of 0.2-0.4 s. The synthetic seismograms computed from
our model are given together with the observed seismograms in Figs. 9-12. In these
computations, the phases corresponding to local reflections are not included. It is seen
that our model explains well the observed amplitude behaviours of PiP2 and PmP phases.
4.3 Modelling for localized reflections
As is stated in Sec. 3, several localized but clear later phases were observed in all
of the record sections (e.g., L1-L4 in Fig. 3). The interpretation for these phases is
rather subjective. For example, they are explained as a focusing effect arising from the
Vol. 41, No. 3, 1993
182 T. Iwasaki et al.
Fig. 13. Modelling for localized reflections. Two possible models for localized
reflections are schematically depicted. (a) Thin low-velocity layer model. (b)
Thin high-velocity layer model.
(a)
(b)
Fig. 14. Synthetic seismograms of S-1 and S-4. Localized reflections are
interpreted by the thin low-velocity layer model as in Fig. 13(a). Locations of
local reflector are shown in Fig. 6(b). These seismograms explain very well the
features of the observed sections in Figs. 9(a) and 12(a). (a) Synthetic
seismograms of S-1. (b) Synthetic seismograms of S-4.
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 183
complex 3-D velocity structure. However, it is impossible to derive such a 3-D structure
only from our experiment. In the present paper, we interpreted that these later phases
had been generated by local reflectors existing under our profile.
The locations of the reflectors were determined from the travel time data (Fig.
6(b)). They are distributed in a depth range of 10-15 km with a horizontal extent of
15-20 km. The large amplitudes of these phases require large velocity changes of
0.3-0.4 km/s at the corresponding boundaries of the reflectors. The velocity structure
shallower than 15 km, however, is well constrained by the travel times of the Pg phase.
In order to explain the large amplitudes of local reflections without affecting the travel
times of the other predominant phases, the thickness of the reflector must be thinner
than 0.5-1 km. If we assume a thick reflector of 3 km, for example, the first arrival times
are not explained well. The thin reflector is also supported by high frequency components
contained in the local reflections (see Fig. 3). Figure 13 shows possible models for the
thin local reflector. In Model A, the reflector is represented by a low velocity zone,
while, in Model B, it is filled with a high velocity material. These two models yield large
amplitude reflections with opposite polarity to each other. Hence, the polarity of the
reflected wave is an important key to understand the physical property of the local
reflector, but difficult to be identified in our seismic data due to the poor S/N ratio.
The Pg phase from S-4, however, is not well explained by a model with a high velocity
reflector because a low velocity region below the reflector yields a shadow zone for this
phase. Therefore, we adopted Model A as a local reflector and computed again the
synthetic seismograms for S-1 and S-4 (Fig. 14). In these examples, the velocities within
the reflectors were set 0.35 km/s lower than those of the surrounding materials. It is
seen that the observed seismograms in Figs. 9(a) and 12(a) are satisfactorily interpreted
by the present modelling.
5. S-Wave Velocity Structure
Record sections of S-waves from S-1 and S-4 are presented in Fig. 15. The S-wave
is generally enhanced by applying the low-pass filter of 6 Hz. In the S-wave record
sections, the reduction velocity is taken to be 3.46 km/s (= 6.0/ã3 km/s) and a time
axis is compressed by a factor of ã3 . This format allows easy comparison between
P-wave and S-wave sections. Namely, if the VP/VS ratio is ã3 in the whole crust and
upper mantle, all of the travel time branches of the S-wave overlay those of the P-wave.
A clear phase with an apparent velocity of 3.45-3.47 km/s is a refracted wave from
the upper crust (Sg phase). The S-wave refracted from the upper mantle (Sn phases) is
observed as a weak first arrival with an apparent velocity of 4.3 km/s at offsets greater
than 170 km from S-1 (Fig. 15(a)). The later phases in these record sections are reflections
from the mid-crustal interface (SiS phase) and the Moho boundary (SmS phase) because
their travel time curves show good correspondence to those of the PiP (PiP2) and PmP
phases in Figs. 9 and 12.
The travel time analysis for these phases indicated that they were radiated as a
S-wave in the vicinity of the source, for which the following two explanations are
plausible. (1) The S-waves were generated at the individual shot point. Explosives are
usually shot in a long (50-60 m) bore hole. Hence, the radiation pattern of the source
Vol. 41, No. 3, 1993
184 T. Iwasaki et al.
(a)
(b)
Fig. 15. Record sections of S-wave data. The reduction velocity is taken as
3.46 km/s (6.0/ã3 km/s) and the time axis is compressed by ã3 relative to
the P-wave time axis. Travel time curves of prominent S-wave phases show
good correspondence to those of P-wave in Figs. 3, 9 and 12. (a) Record section
of S-1. (b) Record section of S-2.
is not spherically symmetric and the S-wave may be yielded at the time of the detonation.
(2) The S-waves were P-S converted waves at a free surface or a very shallow interface
with a large velocity contrast located in the vicinity of the shot point.
The travel times of S-waves in the two cases are nearly identical. Therefore, our
travel time computation was carried out assuming the S-waves were directly radiated
from the source. In the modelling for the S-wave data, we fixed the mid-crustal interfaces
and the Moho determined from the P-wave data and estimated the VP/VS ratio for the
individual crustal layer. The VP/VS ratio in the surface layer was assumed to be 1.73
J. Phys. Earth
Seismic Refraction Study in the Kitakami Region 185
because it was not constrained from our data. The obtained result is summarized in
Fig. 6(b).
The S-wave velocity in the uppermost crust shows a clear lateral change at the
HTB. Namely, the average VP/VS ratio is 1.72-1.73 north of the HTB whereas 1.75-1.76
south of the HTB. The S-wave velocity at the top of the basement is 3.47 and 3.42-
3.45 km/s in the two terranes. It is noted that this lateral velocity change is opposite
in sense to the case of the P-wave velocity (see Sec. 4). The obtained result strongly
indicates the petrological difference between the two geological units. The VP/VS ratio
in the lower crust is 1.73-1.76 although it is less constrained than in the upper crust.
6. Discussion and Conclusions
The Kitakami region is geologically separated into two units by the Hayachine
Tectonic Belt (HTB). The southern part consists of pre-Silurian basements and
Silurian-lower Cretaceous marine sediments, while the northern one is characterized by
a Jurassic accretionary complex. At the Cretaceous time, the heavy granitic intrusion
occurred in both the terranes. A 194-km profile with N-S direction was shot in the
Kitakami region, in 1990. The record sections obtained are characterized by clear first
arrivals (Pg phase refracted from the upper crust) and several later phases (wide-angle
reflected waves from mid-crustal interfaces (PiP phases) and the Moho boundary (PmP
phase)), from which a precise velocity model was constructed.
Our experiment area is covered with a thin surface layer with a velocity of
3.1-5.4 km/s. The thickness of this layer is less than 1 km along the profile. The upper
crustal velocity shows a remarkable regional difference along the profile. The P-wave
velocity at the top of the basement is 5.90 km/s north of the HTB whereas 6.05-6.15 km/s
south of the HTB. Such a velocity difference is recognized down to a depth of 15 km
at which the P-wave velocity is about 6.45 km/s. The high velocity under the southern
Kitakami terrane is in good agreement with those by Matsuzawa (1959) and Minoura
and Hasegawa (1992). According to the map of the upper crustal velocity by Minoura
and Hasegawa (1992), there exists another high velocity region around the northern
edge of our profile. In our model, however, the structure of the upper crust was well
resolved by the dense travel time data. Particularly, the data north of the HTB cannot
be explained by a high velocity of 6.1-6.2 km/s.
The P-wave velocity in a depth range of 15-25 km, which was mainly determined
from the amplitude data, is 6.45-6.65 km. A mid-crustal interface corresponding to a
prominent PiP phase is located at a depth of 25 km in the northern part of the profile,
while 20 km in the southern part. The depth change of this boundary occurs beneath
the HTB. The Moho depth decreases southward from 34 to 32 km. The P-wave velocity
between the mid-crustal interface and the Moho is 6.8-7.0 km/s, which is 0.2-0.4 km/s
higher than the lower crustal velocity by Yoshii and Asano (1972), Zhao et al. (1990),
and Minoura and Hasegawa (1992). If we assume a relatively low velocity of 6.6 km/s
in the lower crust, the predicted amplitudes of PiP phase become much smaller than
the observed ones. Such a model yields a large PmP phase, which also contradicts the
observation. Although the velocity of the uppermost mantle (Pn phase) was not
constrained well by our data, the weak arrivals observed at offsets of 160-194 km suggest
Vol. 41, No. 3, 1993
186 T. Iwasaki et al.
that the Pn velocity is less than 7.7 km/s, This is consistent with the Pn velocities obtained
by Yoshii and Asano (1972) and Zhao et al. (1990).
Besides these clear interfaces, local reflectors are distributed in a depth range of
10-15 km with a horizontal extent of 15-20 km. The thickness of the reflector is very
thin (0.5-1.0 km) and its velocity is 0.3-0.4 km/s lower than that of the surrounding
material.
Clear S-waves observed in our experiment were interpreted as the refraction from
the upper crust (Sg phase) and reflections from the mid-crustal interface (SiS) and the
Moho boundary (SmS phase). The S-wave structure in the upper crust also shows a
remarkable regional difference. The S-wave velocity at the uppermost crust is 3.47 km/s
north of the HTB but 3.42-3.45 km/s south of the HTB. The average VP/VS ratios in
the upper 15-km crust are 1.72-1.73 in the north of the HTB, which is almost comparable
to those by the seismic refraction studies in central Japan (Sasatani et al., 1990; Matsu'ura
et al., 1991). The VP/VS ratio in the southern part of the profile is 1.75-1.76 and
considerably higher than that in the northern part. Such a high value has not been
reported for the upper crust by the previous refraction experiments conducted in the
Japanese Islands. The obtained lateral change in VP/VS ratio represents the petrological
difference between the northern and southern Kitakami terranes. The VP/VS ratio in
the lower crust was determined to be 1.73-1.76. This value is in contrast with the high
VP/VS ratio of 1.77-1.79 in the continental crusts (Perchuc, 1992; Holbrook et al., 1988).
Our experiment recorded weak arrivals with a velocity of about 4.3 km/s, which was
interpreted to be an S-wave refracted from the uppermost mantle (Sn phase).
The obtained crustal structure provided us with a direct seismological evidence
that the Kitakami region is clearly divided into two parts by the HTB. The velocity
difference across the HTB is remarkable and traced down to a depth of 25-30 km. As
stated before, the heavy granitic intrusion took place in the Kitakami region at the
Cretaceous time. Our results strongly indicate that this intrusion was not sufficient to
homogenize the whole crust under the Kitakami region, and the original structural
difference between the northern and southern Kitakami terranes still remains.
The authors wish to express their sincere thanks to members of the Research Group for
Explosion Seismology, particularly, those who participated in the experiment. They also would
like thank to Dr, H. Sato, the University of Tokyo, Profs. M. Ehiro, K. Minoura, Tohoku
University and Prof. K. Okami, Iwate University for their valuable comments on the geology of
the Kitakami region. They are grateful to Prof, M. Matsu'ura for helpful suggestions on the
inversion theory. The experiment was conducted by the Fund of the Special Works of Earthquake
Research Institute, the University of Tokyo, as one of the disciplines of the Japanese Earthquake
Prediction Program.
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Iwaskai et al. 1993 Seismic Refraction Study in the Kitakami Region, Northern Honshu, Japan.txt
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Earth Planets Space,57,727-741,2005 Utility of petrophysical and geophysical data to constrain the subsurface structure of the Kitakami plutons, northeast Japan Shigeo Okuma and Hiroshi Kanaya Geological Survey of Japan,National Institute of Advanced Industrial Science and Technology (AIST), Tsukuba Central 7,Higashi 1-1-1,Tsukuba,Ibaraki 305-8567,Japan (Received December 16, 2004; Revised August 4, 2005; Accepted August 11, 2005) The Geological Survey of Japan (GSJ), AIST has been systematically measuring physical properties of Rock 21 (Petrophysical Database of Basement rocks in Japan for the 21st Century, http://www.aist.go.jp/RIODB/ PB-Rock21/). The petrophysical data has been integrated with aeromagnetic data to determine the subsurface structure of the Himekami and Sobanokami plutons, members of the Kitakami plutonic belt, northeast Japan. For the Himekami pluton, the result of apparent magnetization intensity mapping clearly shows zonation within sub- sub-plutons. The bottom depths of the pluton were estimated by a successive 3-D magnetic modeling to be 1.7 km (North sub-pluton) and 1.5 km below sea level (South sub-pluton), respectively. Magnetic modeling of the Sobanokami pluton, which is partly exposed at the eastern edge of the Ishinomaki Plain, implies that the main part of the pluton extends to the northwest below the surface. Its western edge is cut by the Ishinomaki-wan fault, believed to be the upward extension of the fault plane of the 2003 northern Miyagi earthquake as defined by seismic observations and refection seismic surveys conducted after the earthquake. Integrated magnetic and gravity modeling suggests that the Kakeyama Formation, a Neogene conglomerate deposited in a half-graben along the Ishinomaki-wan fault during Miocene rifting, contains more granitic fragments toward the south than to the north in the Ishinomaki Plain. Key words: Kitakami pluton, petrophysical data, rock magnetic data, magnetic anomaly, gravity anomaly. 1.Introduction magnetic relative to the magnetic granites. This means it is As an example of the use of petrophysical data to inter- straightforward to estimate the dimensions of the Kitakami pret geophysical data, we show here a case study conducted1 plutons by interpretations of the magnetic anomalies in the in the Kitakami Mountains, northeast Japan (Fig. 1), where area. the Early Cretaceous Kitakami Granites intrude Paleozoic The Kitakami Granites are relatively dense (average dry and Mesozoic sedimentary rocks. and wet densities >2.75 × 103 kg/m3) and roughly equiv- The Geological Survey of Japan (GSJ), AIST has a alent to that of Paleozoic and Mesozoic sedimentary rocks, been measuring rock physical data mainly of granitic rocks in Japan under the same laboratory conditions1 lies caused by the granitic rocks, though intensive studies for more than 30 years and has started to construct a were conducted during the initial stage of gravity surveys for their applications to structural interpretations of the ma- Database of Basement rocks in Japan for the 21st Cen- jor plutons in the area (e.g. Kano and Research group of tury, http://www.aist.go.jp/RIODB/PB-Rock21/; Okuma et Granite plutons, Akita Univ., 1978). al., 2003). Based on 391 measurements (Fig. 1), the Ki- In order to understand subsurface structures in the north- takami Granites are the most magnetic (magnetic suscep- western and southwestern edges of the Kitakami Moun- tibility >0.01 SI) on average among pre-Tertiary granites tains, we primarily use magnetic models constrained by the which crop out in the Japanese Islands, whereas the in- rock property measurements and show results of magnetic tensity of Natural Remanent Magnetization (NRM) is rel- interpretations. atively low. Consequently, the Konigsberger (Qn) ratio is low (<0.6) and the effects of NRM can be considered negli- Geologic Setting 2. gible when interpreting magnetic anomalies caused by these The northern Tohoku region is underlain by pre-Tertiary rocks. In contrast Paleozoic and Mesozoic sedimentary basement rocks of the North and South Kitakami belts and rocks surrounding the Kitakami Granites are almost non- Abukuma belt, which are mainly exposed in the Kitakami and Abukuma mountains on the Pacific coast side (Fig. 2; Copyright The Society of Geomagnetism and Earth, Planetary and Space Sci- Oide et al., 1989). The North Kitakami and Abukuma belts comprise the Jurassic accretionary complexes, while the of Japan;TheGeodetic Society ofJapan;The JapaneeSocietyforPlanetary Sci South Kitakami belt is a collided continental block com- ences; TERRAPUB. posed mainly of Paleozoic sedimentary rocks deposited in 727 728 S.OKUMAANDH.KANAYA:SUBSURFACESTRUCTUREOFTHEKITAKAMIPLUTONS PacificOcean 88 Japan Sea A4 SamplingSites 50km is 50 nT. Specification of the original aeromagnetic survey by NEDO is summarized in Table 1. Black dots are rock sampling sites. The area bounded by black solid/broken lines indicates rock sampling areas for the petrophysical database, PB-Rock 21: A1; Kitakami Mountains area, A2; Taiheizan area, A3; Kurikoma-Narugo area, A4; Murakami area. The rectangles bounded by blue solid lines show case study areas for magnetic modeling: 1; Kitakami Mountains area, 2; Himekami Area, 3; 2003 northern Miyagi earthquake area, 4; Sobanokami pluton area. The white line with X and Y shows the location of a magnetic profile (Fig. 10). Blue dots are oil exploration wells: #1; Off-Hachinohe well, #2; Off-Kuji well, #3; Off-Kesennuma well. Table 1. Aeromagnetic : Survey Area Organization Survey Reduced Surface Line Spacing Reference Year (mASL) (km) Tohoku NEDO*1 1981-1982 2,438 3 Okuma (1993) Sengan MITI*2 1978 1800 1-2 Okuma and Suto (1987) (Hachimantai) Abukuma GSJ'3 1970 457 2-3 GSJ(1974) *1 NEDO: New Energy and Industrial Technology Development Organization. *2 MITI : Ministry of Industrial Science and Technology. *3 GSJ: Geological Survey of Japan. a shallow marine environment. These geologic belts are parts (Kanisawa, 1974). The Abukuma belt was also in- believed to have been formed along the Asian continental truded by plutonic rocks during the period from 90 to 100 margin by the end of the Jurassic period and successive in- Ma. The Kitakami Granites represent these plutonic rocks tensive volcanic activity occurred by the oblique subduction and are composed of more than 50 named plutons, which of the Izanagi Plate (Maruyama and Seno, 1986). are classified into six zones (Zone I~VI) based on their Plutonic rocks intruded in the North and South Kitakami mode of occurrence, megascopic features and petrography belts in Early Cretaceous between 110 and 120 Ma and (Katada, 1974; Fig. 3). formed Cretaceousvolcano-plutonic complexesin some The Japan Sea opened in the early and middle Miocene S. OKUMA AND H.KANAYA:SUBSURFACE STRUCTURE OF THE KITAKAMI PLUTONS 729 +++ ceousPlutonicRocks Pacific Ocean North Kitakam Belt) 40°00'N Kitakami Mtns. Japan Sea South F3 Kitakan Belt F4 Abukuma Belt F2 Abukuma AshioBelt Mtns. 50km 14100'E 14200'E Fig. 2. Simplified geologic map of the northern Tohoku Area (modified from Geological Survey of Japan, 1995). Cretaceous plutonic rocks are shown with major fault lines, F1; Hayachine eastern boundary fault, F2; Futaba fault, F3; Hatakawa shear zone, F4; Tanakura fracture zone (after Oide et al. 1989). The rectangle indicates the area of a detailed map of the Kitakami plutons (Fig. 3). between 25 and 15 Ma (Jolivet et al., 1994), causing a drift The uniformity of measurement practices insures repeata- of northern Japan away from the Asian continental margin bility. to roughly its present location. Kanaya and Okuma (2003) have examined the petrophys- ical characteristics of granitic rocks from the northern To- 3. Petrophysical Database and Characteristics of hoku region (Fig. 1) and discussed the general characteris- Petrophysical Data in the Northern Tohoku Re- tics of the data among four sub-divided areas: the Kitakami gion Mountains, Taiheizan, Kurikoma-Narugo and Murakami ar- We have been measuring petrophysical properties of eas. The mean dry density of the four areas varies from 2.66 basement rocks such as granitic and metamorphic rocks in 1 to 2.77 × 103 kg/m3 with the highest values in the Kitakami Japan to better understand the geological and geophysical Mountains (Fig. 4(b)). As for the mean magnetic suscep- structures of the Japanese Islands (Kanaya, 1974; Kanaya tibility, the Kitakami Mountains is also highest among the et al., 1998; Okuma and Kanaya, 1990; Okuma et al., areas (Fig. 4(a)). The Qn ratio of more than 70 percent of 1993; Tanaka and Kanaya, 1986, 1987). Based on the data, granitic rocks is less than 0.4 in this region, implying that we have initiated development of a petrophysical database, NRM is negligible when analyzing magnetic anomalies. called PB-Rock 21 and included the data of basement rocks Comparing dry and wet densities and magnetic suscepti- from the northern Tohoku region in 2003. We will succes- bility of the rocks among the Kitakami Mountains, the val- sively add more data from other areas to cover the whole ues of Zone IV are the highest among the six zones (Fig. 5), Japanese Islands in several years. The database itself con- suggesting a quantitative difference of mafic rocks among sists of petrophysical data such as magnetic susceptibil- the zones (Oide et al., 1989). ity, intensity of NRM, Qn ratio, dry and wet densities and porosity with additional information on lithology, sampling 4. Characteristics of Geophysical Data in the Ki- locations and registration numbers of rock samples at the takami Mountains Geological Museum of the GSJ. Limited measurements of The GSJ conducted regional gravity surveys in the north- Curie temperature and ultrasonic velocity are also available. ern Tohoku region in the 1990's and published a gravity 730 S.OKUMA ANDH.KANAYA:SUBSURFACE STRUCTURE OF THEKITAKAMIPLUTONS (a) Kita amiGranites EarlyCretaceous Volcanic Rocks Ultramafic rocks Mt.Nanas _Sus (SI) Murakam All Area 40N Mt (b) Dry_Density Wet_Density Kitakami Kurikoma- Murakami Dry_Density (x103 kg/m3) All Areas Fig. 4. Petrophysical properties of granitic rocks in the northern To- hoku region. (a) Magnetic properties of granitic rocks in the northern Tohoku region. Mag_Sus: Mean magnetic susceptibility (SI), NRM: 39°N Mean intensity of Natural Remanent Magnetization (A/m), Qn: Mean Konigsberger ratio. The data with Qn ratio greater than 5.0 (18 data) were excluded for averaging. The numbers of rock samples (numbers in parentheses are original number of samples): Kitakami; 373 (391), Taiheizan; 46 (46), Kurikoma-Narugo; 53 (53), Murakami; 86 (86). (b) Densities of granitic rocks in the northern Tohoku region. Dry-Density: Vb VI-a Mean dry density (×103 kg/m3), Wet Density: Mean wet density (× 103 kg/m3). See also Fig. 4(a). Sendai Plain Pacific Ocean map of the Kitakami district (Bouguer anomalies) (Ko- mazawa et al., 1996) as well as adjacent areas. Further- Matsushin more, Komazawa (2000) compiled a nationwide gravity ?口 OjikaPen. map based on these surveys including other agency's data and disclosed gravity data in a CD-ROM. According to a Bouguer gravity map (o = 2.67 × 103 kg/m3) of the Ki- 20 km 0 takami Mountains (Fig. 6), the correlation between the dis- 141E 142PE tribution of the Kitakami Granites and gravity signatures is not obvious. This is supported by the fact that the average Fig. 3.Zonal arrangements of the Kitakami Granites (after Katada, density of the Kitakami Granites and surrounding Paleozoic 1974). I~VI(a/b) indicate zone numbers. Thick solid lines indicate the boundaries of each zone. Kitakami Granites and ultramafic rocks and Early Mesozoic mudstones and sandstones are almost are plotted with abbreviations for the names of each pluton. Kitakami the same (Tanaka and Kanaya, 1986, 1987). An exception plutons: KK; Kuki, TR; Taro, OU; Oura (Zone I), HG; Hashigami, is the Tono pluton, the largest in the Kitakami Mountains: TM; Tenjinmori, TH; Tanohata, MK; Miyako (Zone I), HN; Hiraniwa, an obvious oval gravity low is associated with the pluton. SK; Sakainokami-dake (Zone II), HIM; Hinomiko, HM; Himekami (Zone IV), TN; Tono, KH; Kurihashi, KS; Kesengawa, GY; Goyosan, o seoe na pazaiee (an) i na qn HK; Hitokabe, SM: Senmaya (Zone V), KS; Kinkazan (Zone VIa), the pluton and estimated a depth extent of 10-15 km be- TS; Tabashine, UN; Uchino, OK; Orikabe, HT: Hirota (Zone VIb). low the surface with density decreasing toward the center Ultramafic Rocks: HU; Hayachine Ultramafic Rocks, MU; Miyamori Ultramafic Rocks. #2 and #3 show the locations of oil exploration wells, of the body. This hypothesis for the density distribution is Off-Kuji and Off-Kesennuma. Grey rectangles and triangles with names confirmed by the petrophysical data (Nabetani et al., 1972; denote locations of major towns and mountains, respectively. Tanaka and Kanaya, 1987; Fig. 6). Aeromagnetic surveys have been conducted in the study area and the specifications of data employed in this paper S.OKUMAANDH.KANAYA:SUBSURFACESTRUCTUREOFTHEKITAKAMIPLUTONS 731 (a) 口Or (b) All Are Fig. 5. Petrophysical properties of the Kitakami Granites. (a) Magnetic properties of the Kitakami Granites. I-VI are zone numbers. The numbers of rock samples (numbers in parentheses are original number of samples): I; 50 (54), II; 81 (82), IIl; 34 (35), IV; 26 (26), V; 120 (129), VI; 62 (65). See also Fig. 4(a). (b) Densities of the Kitakami Granites. See also Figs. 4(b) and 5(a). are summarized in Table 1. Aeromagnetic anomalies in the Kitakami Mountains are well correlated to the distribution of the Kitakami Granites and most of the plutons can be recognized as magnetic highs (e.g. Okuma, 1993; Fig. 7). 5. Magnetic Modeling 5.11 Himekami pluton Magnetic models were constructed for several plutons. (FE) The Himekami pluton which crops out about 20 km north- northeast of Morioka City, Iwate Pref. (Fig. 3) is a complex Fig. 6. Bouguer gravity map of the Kitakami Mountains area (after of mafic rocks and their differentiated felsic counterparts Komazawa, 2000). Assumed Bouguer density: 2.67 × 103 kg/m3. and the main pluton is divided into two sub-plutons: the Contour interval is 1 mgal. H and L denote high and low anomalies, South and North plutons (Katada et al., 1991). The North respectively. The rectangles 2 and 4 denote magnetic modeling areas; pluton (A and B in Fig. 8(a) is composed mainly of felsic Himekami and Sobanokami plutons, respectively. A thick solid line with X and Y shows the location of a magnetic profile (Fig. 10). See rocks, whereas the South pluton (C in Fig. 8(a)) is com- also Fig. 3. posed of both mafic and felsic rocks. The North pluton is divided further into three bodies (Katada et al., 1991) but roughly classified into two parts (A and B in Fig. 8(a)) by the contents of SiO2 and rock magnetic data that are de- ites (Fig. 8(b)). Apparent magnetization mapping (Okuma scribed later. et al., 1994; Nakatsuka, 1995) was conducted assuming the A long-wavelength dipolar magnetic anomaly indicat- top and bottom extents of the magnetic structure correspond ing a magnetization vector in the Earth's magnetic field to the terrain surface and 16 km below sea level, respec- lies over the pluton. Within this anomaly is a short- tively on the basis of a magnetic model which corresponds wavelength dipolar magnetic anomaly with the opposite po- to the buried Kitakami plutons along the Pacific coastline of larity, indicating the heterogeneity of the Himekami Gran- northern Tohoku (Finn, 1994). The resultant apparent mag- 732 S.OKUMAANDH.KANAYA:SUBSURFACESTRUCTUREOFTHEKITAKAMI PLUTONS two parts: A and B show the mean values of 0.38 A/m amiGranites SI), respectively. fragments include igneous rocks, limestone, and chert rock magnetic properties of the pluton (Table 2) and are closely associated with petrography. The North pluton is a zoned pluton of quartz monzonite (SiO2 content: 60-65%), granite (63-65%) and granodiorite (67-70%) (Katada et al., 1991). The most felsic part of the North pluton, A, is com-- posed of granodiorite and shows the lowest magnetic sus- ceptibility of 1.72 × 10-2 SI. Whereas, B comprises quartz monzonite and granite, with a moderate magnetic suscep- tibility of 4.79 × 10-2 SI. The South pluton, C is com- prised mainly of quartz monzonite (SiO2 content: 55-58%) and quartz monzodiorite (52-54%) as felsic rocks (Katada et al., 1991), showing the largest magnetic susceptibility of 9.43 × 10-2 SI which is almost double that from B and five times greater than that of A. The South pluton is further divided into western and eastern parts based on the horn- blende content, and the eastern part contains some mafic rocks such as monzogabbro (SiO2 content: 49%) (Katada et al., 1991). The highest magnetic susceptibility of C was measured on specimens sampled from gabbroic outcrops in the study area. There is an obvious difference in appar- ent magnetization intensity between the western and east- ern parts of area C: the western part is less magnetic (<1.0 A/m (k = 2.65 × 10-2 SI) and the central - eastern part is more magnetic (~3.0 A/m (k = 8.0 x 10-2 SI)) (Fig. 8(c)). The apparent magnetization intensity map implies that the felsic rocks in the eastern part are underlain by mafic rocks. of the apparent magnetization intensity mapping and rock magnetic measurements are well correlated, the actual rock magnetic properties tend to be higher than the apparent magnetization intensities, suggesting that the bottom depth cifidocean of the pluton is shallower than the initial assumption (16 km BSL) of the mapping. Thus, we tried to estimate the bottom depth of the pluton. We assumed an ensemble of vertical prisms with the horizontal boundaries deduced from a geo- logic map (Katada et al., 1991) for the North pluton (A and B) and an apparent magnetization intensity map (Fig. 8(c)) 2okm for the South pluton (C) as the magnetic structure, which is L41°E A42PE subject to the magnetic anomalies. We allowed variations of uniform bottom depths of vertical prisms for the North and Fig. 7. Reduction to the pole anomaly map of the Kitakami Mountains South plutons individually, but fixed the tops to the terrain area (after Okuma, 1993). Contour interval is 25 nT. Solid and broken height. Synthetic magnetic anomalies (Fig. 8(d)), which contours indicate positive and negative values, respectively. H and L denote high and low anomalies, respectively. See also Fig. 6. best fit the observed residual anomalies (Fig. 8(b)), were calculated by trial and error, changing the bottom depths in are 0.48 and 0.40, much higher than those of the other netization intensity map (Fig. 8(c) clearly shows the mag- netization highs lie over the Himekami pluton and anoma- Uyeda, 1967) to judge how synthetic anomalies (G) fit well lies are caused by the pluton. The map also shows a big to the observed anomalies (F): the larger the ratio, the difference in magnetic properties between the North and better the fit. South plutons. The calculated magnetization intensities of Goodness-ratio (r) can be written by the South pluton reach 3.0 A/m (k = 8.0 × 10-2 SI, when F = 47, 500 nT) in some parts with the mean value of 1.24 F; A/m (k = 3.28 x 10-2 SI), whereas those of the North plu- (1) ton range from 0.25 to 1.0 A/m (k = 0.66-2.65 × 10-2 SI). Furthermore, the North pluton can be classified into Gi S.OKUMAANDH.KANAYA:SUBSURFACE STRUCTURE OF THEKITAKAMI PLUTONS 733 Topography Total MagneticIntensity P P2 P1 14110'E +141:Y5' 5km P2 5km P2 (a) (b) MagnetizationIntensity Synthetic Anomaly P2' (nT 39°55'N 39° 55' N P 39° 45' N 41°10'E 141915E 14110'E 14115'E 5km 5km P2 (c) (d) Fig. 8. Magnetic modeling of the Himekami pluton. (a) Topographic map of study area. Contour interval is 100 m. Thick grey lines indicate geologic boundaries of the pluton (simplified from Katada et al., 199i). A and B denote less magnetic and magnetic parts of the North pluton, respectively. C is the South pluton. Black dots indicate rock sampling sites. (b) Total magnetic intensity anomaly map of study area (after Okuma and Suto, 1987). Contour interval is 25 nT. Specification of the original aeromagnetic survey by MITI is summarized in Table 1. Thin black solid lines indicate flight lines, spaced 1-2 km apart. See also (a). (c) Apparent magnetization intensity map of study area. Contour interval is 25 × 10-2 A/m. Thick white lines indicate geologic boundaries of the pluton estimated by the apparent magnetization mapping. See also (a). (d) Synthetic magnetic anomaly that best fits the observed one (b)). Contour interval is 25 nT. See also (c). 734 S.OKUMAANDH.KANAYA:SUBSURFACESTRUCTURE OF THEKITAKAMIPLUTONS 5.0 600.0 (nT） 4.0 P1 TMI (Observed) P1' 400.0 TMI (Calculated) (E) 200.0 0.0 200.0 Topography 0.0 ·B' 4.79 x 10-2 SI 600.0 2.0 800.0 4.0 6.0 8.0 10.0 12.0 14.0 Distance (km) (a)E-WProfile 600.0 nT P2 + TMI (Observed) P2' TMI (Calculated) (S) (N) 200.0 0.0 9.43x10-2S Topography 200.0 0.0 600.0 2.0 1.72 x 10-2 SI 0.0 2.0 4.0 6.0 8.0 10.0 12.0 14.016.0 18.0 20.022.0 Distance (km) (b) N-S Profile Fig. 9. Cross-sections of the magnetic model of the Himekami pluton. (a) E-W profile: East-West cross-section of an optimum magnetic model (Fig. 8(d)) which best fits the observed total magnetic intensity anomaly (Fig. 8(b)). Vertical exaggeration is 2.0. Numerals stand for magnetic susceptibilities employed for the magnetic modeling. See also Fig. 8(a). (b) N-S profile: North-South cross-section of an optimum magnetic model (Fig. 8(d) which best fits the observed total magnetic intensity anomaly (Fig. 8(b)). See also Fig. 9(a). Table 2. Rock magnetic properties of the Himekami pluton. Body MagneticSusceptibility NRM Intensity Qn* Number Number Average Average ratio of Sampling of Rock (Standard Dev.) (Standard Dev.) Sites Specimens (SI) (A/m) 1.72 x 10-2 2.09 x 10- A 0.32 2 6 (4.09 x 10-4) (6.42 x 10-2) 4.79 x 10-2 7.95 x 10-1 B 0.43 (8.40 x 10-3) 2 6 (2.51 x 10²) 9.43 x 10-2 1.93 x 100 C 0.54 3 21 (4.38 x 10-2) (8.51 x 10-) *: The Earth's magnetic field of 47,500 nT was assigned for the calculation of Qn ratio. 735 400.0 16.0 200.0- 12.0 Observed RTP Calculated RTP 0.0 (1u) 8.0 4.C coastline Ano Himekami Sakainokami Miyako Taro Topography Pole PacificOcean 400.0 0.0 (3.37x10-) (2.84x10-2) (A: 1.72102,B: .79x102, C:9.4x102 600.0 -4.0- (2.14x10°²) Redi Paleozoic / Cretaceus Sediments -800.0 -8.0 1000.0 12.0 1200.0 0.0 20.0 30.0 40.0 50.0 60.0 70.0 80.0 Distance (km) Fig. 10. Schematic cross-section of the magnetic structure along the profle X-Y. See the location of the profile in Fig. 1. A line and plus signs in the upper part of the figure indicate observed and calculated reduction to the pole anomalies (RTP), respectively. Vertical exaggeration is 2.5 for the structure. Grey areas denote cross-sections of the Kitakami plutons; Himekami, Sakainokami, Miyako and Taro from left to right, whose structures were estimated by individual magnetic modeling as isolated anomalies without regional trends. Numbers in parentheses denote averaged magnetic susceptibilities (SI) employed for magnetic modeling. Thick broken lines show the bottom depths of each pluton estimated by an MT modeling (Ogawa, 1992). Assuming induced magnetization and the terrain higher the Kitakami plutons, called the Sobanokami pluton ex- than 750 m above sea level is not magnetized due to weath- ists (Fig. 13(b)). This area corresponds to the southwestern ering and does not contribute to the magnetic anomalies, edge of the Kitakami Mountains. The center of the area is optimum bottom depths of the vertical prisms are 1.7 km covered by alluvium and is bounded by Miocene/Pliocene and 1.5 km below sea level for the North and South plu- sedimentary rock to the west and by Triassic sedimentary tons, respectively (Fig. 9). This corresponds to the largest rocks to the east. goodness-ratio of 4.90. This result is a good example that Magnetic modeling was conducted by trial and error cal- implies bottom depths of satellite plutons of the Kitakami culation along the WSW-ENE and NNW-SSE directions Granites are not always so thick (Fig. 10). (Fig. 13(d), assuming an ellipsoid as the magnetic source 5.2 2003 Northern Miyagi Earthquake area (Clark et al., 1986). When using an average magnetic sus- The second example is a study over the 2003 Northern ceptibility of the Sobanokami pluton from the petrophys- Miyagi Earthquake area (Fig. 11) where a series of strong ical database, it is difficult to find a suitable solution for earthquakes (JMA intensities >5) occurred during the pe- the analysis because of low magnetic susceptibility levels. riod from July 26-28 in 2003. In this area, the Asahiyama The average magnetic susceptibility of the Sobanokami plu- Flexure has been known as an active fault (Research Group ton (3.00 x 10-3 SI) calculated from previous petrophysi- for Active Faults of Japan, 1991) but no surface faults cal measurements is conspicuously low compared to that caused by the recent earthquakes were observed. Seismic of the Kitakami Granites (2.69 x 10-2 SI). An additional activity including fore- and aftershocks (most of them <10 sampling of the Sobanokami pluton has been conducted re- km BSL) occurred in an inclined plane with a slope of about cently and implies that the previous petrophysical measure- fifty degrees westward from the ground and the upward ex- ments were collected from outcrops close to contacts with tension of the activity area reaches ground level 5 km east sedimentary rocks. Magnetic susceptibility measurements of the Asahiyama Flexure (Okada et al., 2003). In this of specimens from a newly-found outcrop of the pluton av- area, the northwestern extension of the Ishinonmaki-wan erage 3.36 x 10-2 SI (Table 3) and are almost equal to Fault was thought to extend along the east flank of Sue Hill that of Zone VI (3.31 × 10-2 SI). The average intensity of (Nakamura, 1990; Figs. 11, 12 and 13(a)) but has not been NRM and Qn ratio of the specimens from the outcrops are treated as a serious active fault that causes disastrous earth- 2.25 × 10-1 (A/m) and 0.18, respectively. In this modeling, quakes. parameters of the optimum model are summarized in Table A magnetic high occurs east of Sue Hill, north of Ishi- 4: the width (WSW-ENE), length (NNW-SSE) and height nomaki City (Fig. 13(c)), where small outcrops of one of of the ellipsoid are 4000 m, 6000 m and 1650 m, respec- 736 S.OKUMAANDH.KANAYA:SUBSURFACESTRUCTUREOFTHEKITAKAMIPLUTONS Fault IshinomakiBay 5km Fig. 11. Reduction to the pole anomaly map of the 2003 northern Miyagi earthquake area. Contour interval is 25 nT. This map was created from the total magnetic intensity anomaly map of the area (GSJ, 1974; Table 1). The rectangle bounded by solid line indicates the Sobanokami pluton area (Fig. 13). Line Q1-Q1' and Q2-Q2' show the locations of magnetic modeling profiles with an assumption of an ellipsoid model (Fig. 13). Line R1-R1' and R2-R2' show the locations of integrated gravity and magnetic 2-D modeling profiles (Fig. 15). See also Fig. 1. 5km Fig. 12. Bouguer gravity map of the 2003 northern Miyagi earthquake area. Assumed Bouguer density: 2.67 × 103 kg/m3. Cont This map was created from the gravity database by the GSJ (Komazawa, 2000). See also Fig. 11. 737 yob omaki-wanFault omaki-wanFault TS Sobanokami Sobanokami Granites Granites TS 8°28'N Hill 1km 8°26'N MS TS 141°15'E 41°21'E 141°15'E 141°18'E 141°21'E (a) (b) 01 141°21'E 141°21'E (c) (d) Fig. 13. Magnetic modeling of the Sobanokami pluton area. (a) Topographic map of study area. Contour interval is 10 m. Thick black broken and solid lines indicate faults; Ishinomaki-wan fault (Nakamura, 1990) and Jyobonsan-nishi fault (Research Group for Active Faults of Japan, 1991) from west to east. Thick red solid lines show the outcrops of the Sobanokami pluton. Black dots and a blue open circle denote previous and new rock sampling sites, respectively. (b) Simplified geologic map of study area (modified from Takizawa et al., 1992). TS: Triassic sedimentary rocks, MS: Miocene sedimentary rocks, PS: Pliocene sedimentary rocks, no label: Alluvium. See also Fig. 13(a). (c) Reduction to the pole anomaly map of study area. Contour interval is 10 nT. See also Fig. 13(a). (d) Plan-view of a magnetic model. Crosses with Q1-Q1' and Q2-Q2' indicate the locations of the WSW-ENE and SSE-NNW modeling profiles in Fig. 14, respectively. Light blue lines denote the wire frame of a magnetic ellipsoid. See also Fig. 13(c). Table 3. Rock properties of the Sobanokami pluton. Magnetic NRM Qn* Density Number Number Susceptibility ratio of Sampling of Rock Intensity Inc. Dec. Average Dry Sites Specimens Average (deg.) (deg.) (Standard Dev.) (Standard Dev.) (Wet) (SI) (A/m) (x10°kg/m) 3.36 x 10-2 2.25 x 10- 62.6 -30.1 0.18 2.76 2 24 (3.84 x 10-) (1.05 x 10-) (2.77) *: The Earth's magnetic field of 47,500 nT was assigned for the calculation of Qn ratio. tively, and the top depth of the ellipsoid is just beneath the count NRM of the pluton, the thickness of the ellipsoid is ground and bottom depth ranges from 800 to 2000 m with estimated to be 1200 m. a thickness of 1650 m (Fig. 14). NRM was ignored in this modeling. As the NRM direction of the Sobanokami pluton Discussions 6. is close to that of the present Earth's magnetic field (Table 6.1 Himekami pluton 3), composite magnetic susceptibility can be calculated to Three-dimensional magnetic modeling of the Himekami be about 4 × 10-2 SI by summation. Thus, if we take into pluton (this study) and similar modeling of Sakainokami, 738 S.OKUMAANDH.KANAYA:SUBSURFACESTRUCTURE OF THEKITAKAMIPLUTONS 400 Observed Anomaly 400 Observed Anomaly Calculated Anomaly Calculated Anomaly 300 (1u) 300 RTP 200 RTP 200 100 100 Q1 Distance (m) Q1' Q2 Distance (m) Q2 1000 2000 3000 4000 5000 6000 7000 -1000 1000 2000 3000 4000 5000 6000 7000 8000 9000 -2000-1000 1000 1000 Ishinomaki-wanFault Observed Planet ExtensionofJyobonsan-nishiFault + ObservedPlane :Topography.. C 0 E -1000 1000 -2000 -2000 -3000 -3000 Az= 62.3deg Az=-27.7deg -4000 4000 (a) (b) Fig. 14. Cross-sections of the magnetic model for the Sobanokami pluton area. (a) WsW-ENE cross-section (Q1-Q1'). See the location of the profile in Fig. 13(d). Crosses and a line in the upper part of the figure denote observed and synthetic reduction to the pole anomalies, respectively. A wire frame structure shows a magnetic ellipsoid. The location of the Ishinomaki-wan fault was indicated with a solid arrow and broken line. (b) SSE-NNW cross-section (Q2-Q2'). The estimated extension of the Jyobonsan fault was indicated with a solid arrow. See also Fig. 14(a). Table 4. Model parameters of the Sobanokami pluton. Profile A* B* C* Strike Dip* Plunge* Magnetic Susceptibility (m) (m) (m) (deg.) (deg.) (deg.) (SI) A-A' 4000 6000 1650 -27.7 10.0 8.0 3.36 x 10-2 A, B, C: Width, length and height of an ellipsoid along three semi-axes. Strike: Azimuth of plunge of the axis along the width from the east, Dip: Plunge of the axis (positive direction) along the width from the horizontal plane, Plunge: Plunge of the axis along the length from the plane containing axes along the width and length. Miyako and Taro plutons (in preparation) in the north Ki- South pluton (C) is relatively large while those from the takami Mountains constrained with magnetic susceptibili- North pluton are small (Table 2). This is because C com- ties averaged from rock magnetic data reveal the subsurface prises felsic and mafic parts (Katada et al., 1991). The av- structure of the plutons. The results show that the bottom eraged magnetic susceptibilities of felsic and mafic parts are depths of these plutons are 1.5~1.7 km, 0.35 km, 0.5 km about 6.34 × 10-2 (SI) and 1.56 × 10-1 (SI), respectively. and 3.0 km below sea level, respectively (Fig. 10). If we adopt the MT result and assume the bottom depth of In this area, a magnetotelluric (MT) survey was con- C is 3 km below sea level, an optimal magnetic suscepti- ducted along a E-W profile and the result of a 2-D inver- bility of C is 7.80 × 10-2 (SI). In this case, the ratio of sion suggested the bottom depths of the resistive Himekami felsic to mafic parts can be estimated to be approximately pluton (100-1000 Ω2·m), Sakainokami and Miyako plutons 5:1, whereas the proportion of mafic parts is estimated to be (both >1000 Ω·m) are 3 km, 3 km and 15 km below sea less than 5 percent according to the geologic map (Katada level, respectively (Ogawa, 1992; Fig. 10). The relative et al., 1991). This implies the felsic parts of C are underlain depth variations in both magnetic and resistivity models are by more mafic parts than observed at the surface. similar, although there is a discrepancy in the bottom depth The accuracy of MT modeling is another cause for the of the plutons between the magnetic and MT results. This discrepancy. Three-dimensional structures of the plutons discrepancy might be caused partly by the less accurate as- should also be taken into account for MT modeling, because sumption of magnetic susceptibility for magnetic modeling. dimensions of the satellite plutons are not big enough to as- The reliability of depth estimation by magnetic modeling sume a 2-D model in some cases (e.g. Sakainokami pluton: depends chiefy on determination of proper rock magnetic 3 x 5 km). properties. As for the Himekami pluton, the standard de- Furthermore, uncertainty remains in the analyses of mag- viation of magnetic susceptibility of rock samples from the netic anomalies on the coast of the Kitakami Mountains. As S. OKUMA AND H. KANAYA: SUBSURFACE STRUCTURE OF THE KITAKAMI PLUTONS 739 shown in a magnetic profile (Fig. 10),a regional magnetic rock magnetic measurements are necessary to clarify this high is superimposed on local magnetic highs close to the problem. coast.Before the magnetic modeling of satellite plutons,6.2 2003 Northern Miyagi Earthquake area these regional trends had been removed, and their effect is In the 2003 northern Miyagi earthquake area, steep grav- not included in the result of the modeling at this stage. In ity gradients (Fig. 12) indicate a major boundary between these areas, composite magnetic modeling considering 2-D the Kitakami massif and Tertiary-Quaternary sedimentary and 3-D structures is necessary. rocks (Komazawa et al., 1996). However, it has not been On the Pacific coast of the Kitakami massif, distinctive thought to be associated with active faults which might linear N-S positive magnetic anomalies had been observed cause severe earthquakes due to the cover of Quaternary from the westerm part of Hokkaido and named Kitakami- sediments. Ishikari Magnetic Belt (Ogawa and Suyama, 1976; Oshima Umino et al. (2003) conducted a temporal seismic ob- et al., 1975; Fig. 1). Segawa and Furuta (1978) conducted servation right after the 2003 northern Miyagi earthquake an analysis of these positive magnetic anomalies along the (M6.4) and showed that the fault plane is a curved surface: latitude of 41 degrees N and estimated a volcanic structure the northern part of the plane dips to the west at an angle at a depth of 5 km below sea level with a thickness of 10 of ~50 degrees and the southern part to the northwest at km, assuming a northwestward declination of magnetiza- ~40 degrees.As the upward extension of the depth dis- tion although a low Qn ratio less than 0.5ofvolcanic core tribution of aftershocks coincides with surface trace of the samples from the bottom of two drill holes in the anomaly Ishinomaki-wan fault, they thought the rupture of the M6.4 area in Hokkaido was observed. earthquake took place along the fault at a depth of 3-12 km. Finn (1994) attributed the linear positive anomalies to the On the other hand, reflection seismic surveys were con- Kitakami Granites on the basis of the existence of the Ki- ducted after the earthquake to reveal the subsurface struc- takami Granite at the bottom of a drill hole, off-Kesennuma ture of the area and the extension of a buried fault associated (#3 in Fig. 1; Japan Natural Gas Association and Japan Off- with the recent earthquakes is assumed to exist along the shore Petroleum Development Association,1986),with K-east flank of Sue Hill (e.g. Sato et al., 2004; Yokokura et al., Ar ages of 121±4 and 111±4 for biotite and K-feldspar, 2004). Sato et al. (2004) showed that Miocene sediments of respectively (Shibata, 1986). She also set the bottom depth the rift and post-rift stages such as the Kakeyama Formation of the batholiths around 16 km below sea level,referring to and Pliocene sediments are underlain by Pre-Tertiary base- a resistivity cross-section by an MT survey (Ogawa, 1992) ment to the west of the Ishinomaki-wan fault. and a refraction seismic survey (Iwasaki et al., 1994). The Kakeyama Formation is a basal conglomerate com- Although the positive linear magnetic anomalies on the posed of cobbles and boulders which were supplied from Pacific coast of the northern Kitakami Mountains seem to pre-Tertiary rocks such as shale, sandstone, granitic rocks have a close relationship with the distribution of the Ki- and chert (Takizawa et al., 1992).It is assumed that the takami Granites,we should be careful in examining the formation was deposited in a half-graben bounded by the origin of magnetic anomalies in the area.Actually, the Ishinomaki-wan fault to the east during early Miocene age Harachiyama Formation composed of felsic, intermediate (Sato et al., 2004). The formation was uplifted by reverse of the northern Kitakami Mountains close to the Kitakami the eastern edge of the formation crops out west of the plutons such as in Zone I and II (Figs. 3 and 7). The fault (Nakamura, 1992). Reflection seismic surveys (Sato Harachiyama Formation was formed in the Early Creta- et al., 2004; Yokokura et al., 2004) revealed that the main ceous just before the intrusion of the Kitakami Granites and part of the formation is overlain by Pliocene and Pleis- is believed to form a Cretaceous volcano-plutonic complex tocene sediments and some dips to the west are about 30 with the Kitakami plutons (Kanisawa, 1974). Felsic parts degrees. A portion of the magnetic high which corresponds of the Harachiyama Formation do not show strong mag- to the Sobanokami pluton extends further to the southeast netic properties (Otofuji et al.,1997) but mafic-intermediate (Fig. 11). It might support the idea of the subsurface exis- parts of the formation might cause magnetic anomalies be- tence of conglomerates which include granitic rocks. How- cause of magnetic susceptibilities as high as the Kitakami ever, the gradient of the Bouguer gravity (Fig. 12) is the Granites (in preparation）. Cretaceous mafic volcanic rocks steepest west of the fault not right on the fault contact, im- were retrieved as cuttings from the bottom of the drill hole plying the density of the Kakeyama Formation which abuts Off-Hachinohe (#1 in Fig. 1; Japan Natural Gas Association on the fault contact is comparable with that of Paleozoic and and Japan Offshore Petroleum Development Association, Cretaceous sedimentary rocks. 1992) just below the eastern magnetic high of the Kitakami- Integrated 2-D modeling of magnetic and gravity anoma- Ishikari Magnetic Belt and might correspond to a northern lies (e.g. Talwani, 1965) was additionally conducted to ad- extension of the Harachiyama Formation. dress the problems raised above. The Sobanokami pluton Furthermore, many granitic rocks of the Kitakami plu- was not imaged on reflection seismic and gravity cross- tons are closely associated with gabbroic rocks (Kanisawa sections (Sato et al., 2004), because there is no density and Katada, 1988). Although gabbroic rocks of Zone I, contrast between the pluton and its surrounding Paleo- II and V occupy smaller volumes than the rocks of Zone zoic/Mesozoic sedimentary rocks. In contrast, the pluton I, IV and VIb, the large amplitude of magnetic anomalies was imaged successfully by magnetic modeling in the pre- (Figs. 1 and 7) suggests the granitic rocks are underlain by vious chapter.Thus, the magnetic structure modeled for 740 S.OKUMA ANDH.KANAYA:SUBSURFACESTRUCTURE OFTHEKITAKAMIPLUTONS 200.0 100.0 二 =Observed, Calculated, # 0.0 -100.0 三 -200.0 140.0 =Observed, =Calculated, 120.0 E 100.0 G 80.0 ...... Late Miocene Post Rit Sediment 1 Late Miocene Post Rit Sediment 2 Pliocene/Quaternary 0=-670,S=0.025 A.F. D=-670,S=0 R1 1.F. R1 0.0 +Sobanokami Pluton 1.0 Mioce Kakeyam +D=90,S=0.0336-+ 2.0 170. Srmatio 3.0 D=90,S=0 4.0 XXX 0.0 5.0 10.0 15.0 Distance (km) Fig. 15. Cross-section of the subsurface structure of the 2003 northern Miyagi earthquake area estimated by integrated gravity and magnetic modeling. See the location of the profile R1-R1' in Figs. 11 and 12. The profile was set along the seismic reflection study (Sato et al., 2004). D: Relative has finite lengths of 3 and 5 km in the plus and minus strike directions, respectively while other blocks have infinite lengths. I.F. and A.F. indicate 2004). cross-section (Fig. 15). Another magnetic structure was ton and other satellite plutons are apparently shallower than modeled on and around the Asahiyama Flexure, represent- those estimates from 2-D MT modeling. Magnetic model- ing Miocene volcanic rocks. The density of the Sobanokami ing in the 2003 northern Miyagi earthquake area determines pluton and Paleozoic/Mesozoic sedimentary rocks was as- the subsurface extension of the Sobanokami pluton which sumed to be 2.76 x 103 kg/m3 after the rock property mea- crops out partly at the southwestern edge of the Kitakami surements (Table 3) and is almost equal to the averaged massif and is bounded by the Ishinomaki-wan fault to the value of the Kitakami Granites (2.77 x 103 kg/m3) (Fig. 5). west. Integrated magnetic and gravity modeling suggested The cross-section of the profile R1-R1' (Fig. 15(a)) along granitic fragments from the Sobanokami pluton included in the reflection survey line indicates the Kakeyama Formation the Kakeyama Formation increase toward the south. is relatively dense (2.50 x 103 kg/m?) and slightly magnetic (magnetic susceptibility: 7.0 × 10-3 SI). Furthermore, if we Acknowledgments. We are indebted to Tadashi Nakatsuka for take a cross-section of the profile R2-R2' along the south- the geophysical processing and mapping programs (Nakatsuka, 2003) to plot almost all the maps in this paper. We are grateful western extension of a magnetic high (Fig. 11), this portion to Yasuaki Murata for providing a program, JKGLIB ver. 1 to plot of the formation is analyzed to be more magnetic (magnetic geologic maps of northern Tohoku. We would like to thank two susceptibility: 1.0 × 10-2 SI) than the model of Fig. 15. referees, Carol Finn (US Geological Survey) and Mark Pilking- This suggests the southern part of the Kakeyama Formation ton (Geological Survey of Canada) for their helpful comments to improve the manuscript. contains more granitic fragments as conglomerate than the northern part, probably because of its closeness to the plu- ton's exposure. References Clark, D. A., S. J. Saul, and D. W. Emerson, Magnetic and gravity anoma- 7. 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Ogawa, Y, Deep crustal resistivity structure revealed by wideband Ohtaki, Seismic reflection surveys around the southern part of hypocen- magnetotelluricsTohoku and Hokkaido region, Ph.D. Thesis, Univ. tral area of 2003 Miyagi-ken Hokubu earthquake (part 1), Abstracts of of Tokyo, 1992. the 2004 Japan Earth and Planetary Science Joint Meeting, S053-017, Oide, K., H. Nakagawa, and S. Kanisawa, Regional Geology of Japan, Part 2004. 2 Tohoku, Kyoritsu Shuppan, 338 pp., 1989 (in Japanese). Okada, T., N. Umino, and A. Hasegawa, Rupture proces of the July formed, under the thermal effect of the subducting S. Okuma (e-mail: s.okuma@ aist.go.jp) and H. Kanaya double-difference hypocenter locations, Earth Planets Space, 55, 741-
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Okuma & Kanaya 2005 Utility of petrophysical and geophysical data to constrain the subsurface structure the Kitakmai plutons, northeast Japan.txt
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