The Higo metamorphic complex in Kyushu, Japan as the fragment of Permo–Triassic metamorphic complexes in East Asia

Gondwana Research 9 (2006) 152 – 166 www.elsevier.com/locate/gr

The Higo metamorphic complex in Kyushu, Japan as the fragment of Permo–Triassic metamorphic complexes in East Asia Y. Osanai a,*, M. Owada b, A. Kamei c, T. Hamamoto d, H. Kagami e, T. Toyoshima e, N. Nakano a, T.N. Nam f a

Division of Evolution of Earth Environments, Graduate School of Social and Cultural Studies, Kyushu University, Fukuoka 810-8560, Japan b Department of Earth Sciences, Yamaguchi University, Yamaguchi 753-8512, Japan c Department of Geosciences, Shimane University, Matsue 690-8504, Japan d Dia Consultant Co. Ltd, Nagoya 456-0002, Japan e Graduate School of Natural Science and Technology, Niigata University, Niigata 950-8281, Japan f Department of Geosciences, Hue University of Science, Hue, Vietnam Received 12 November 2004; accepted 29 June 2005 Available online 9 January 2006

Abstract The Higo terrane in west-central Kyushu Island, southwest Japan consists from north to south of the Manotani, Higo and Ryuhozan metamorphic complexes, which are intruded by the Higo plutonic complex (Miyanohara tonalite and Shiraishino granodiorite). The Higo and Manotani metamorphic complexes indicate an imbricate crustal section in which a sequence of metamorphic rocks with increasing metamorphic grade from high (northern part) to low (southern part) structural levels is exposed. The metamorphic rocks in these complexes can be divided into five metamorphic zones (zone A to zone E) from top to base (i.e., from north to south) on the basis of mineral parageneses of pelitic rocks. Greenschist-facies mineral assemblages in zone A (the Manotani metamorphic complex) give way to amphibolitefacies assemblages in zones B, C and D, which in turn are replaced by granulite-facies assemblages in zone E of the Higo metamorphic complex. The highest-grade part of the complex (zone E) indicates peak P – T conditions of ca. 720 MPa and ca. 870 -C. In addition highly aluminous Sprbearing granulites and related high-temperature metamorphic rocks occur as blocks in peridotite intrusions and show UHT-metamorphic conditions of ca. 900 MPa and ca. 950 -C. The prograde and retrograde P – T evolution paths of the Higo and Manotani metamorphic complexes are estimated using reaction textures, mineral inclusion analyses and mineral chemistries, especially in zones A and D, which show a clockwise P – T path from Lws-including Pmp – Act field to Act – Chl – Epi field in zone A and St – Ky field to And field through Sil field in zone D. The Higo metamorphic complex has been traditionally considered to be the western-end of the Ryoke metamorphic belt in the Japanese Islands or part of the Kurosegawa – Paleo Ryoke terrane in south-west Japan. However, recent detailed studies including Permo – Triassic age (ca. 250 Ma) determinations from this complex indicate a close relationship with the high-grade metamorphic terranes in eastern-most Asia (e.g., north Dabie terrane) with similar metamorphic and igneous characteristics, protolith assembly, and metamorphic and igneous ages. The north Dabie high-grade terrane as a collisional metamorphic zone between the North China and the South China cratons could be extended to the N-NE along the transcurrent fault (Tan-Lu Fault) as the Sulu belt in Shandong Peninsula and the Imjingang belt in Korean Peninsula. The Higo and Manotani metamorphic complexes as well as the Hida – Oki terrane in Japan would also have belonged to this type of collisional terrane and then experienced a top-to-the-south displacement with forming a regional nappe structure before the intrusion of younger Shiraishino granodiorite (ca. 120 Ma). D 2005 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. Keywords: Higo metamorphic complex; P – T evolution; Permo – Triassic collision metamorphism; Regional nappe; Tectonics of East Asia

1. Introduction Pre-Tertiary geological units in the Japanese Islands have been regarded as fragments of eastern margin of the Asian * Corresponding author. E-mail address: [email protected] (Y. Osanai).

continent by many workers and these fragments were thought to have split from the Asian continent during Miocene times by the opening of the Japan Sea (East Sea) (e.g., Isozaki and Maruyama, 1991; Tazawa, 2004). The Higo terrane in this study is located in the western part of central Kyushu Island, south-west Japan, and extends for about 40 km in E –W length and has a maximum N – S width of

1342-937X/$ - see front matter D 2005 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2005.06.008

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20 km (Fig. 1). This terrane consists of the Permian Mizukoshi Formation, the Cretaceous Mifune Group, and three metamorphic complexes and is intruded by the Higo plutonic complex. Three distinct metamorphic complexes have been identified and named as the Manotani metamorphic complex, the Higo metamorphic complex and the Ryuhozan metamorphic complex (e.g., Karakida, 1992), which are distributed in northern, central and southern part of the Higo terrane, respectively. The Higo plutonic complex intruded into the Higo and Ryuhozan metamorphic complexes. The Higo metamorphic complex is one of the most widespread high-grade (high-temperature and low-pressure: HT/LP) metamorphic complexes in western Japan and had been identified as the western extension of high-grade ‘‘Ryoke’’ metamorphic belt exposed in the main portion of Japanese islands (Honshu island) (e.g., Yamamoto, 1962; Tsuji, 1967; Yamamoto, 1983). Furthermore, the Higo metamorphic complex has been considered to be part of the polyphase – metamorphosed ‘‘paleo-Ryoke’’ belt, which extends to north-eastern Japan, as deduced from the original rock compositions, metamorphic ages and metamorphic processes (e.g., Karakida et al., 1989; Hayama, 1991; Hara et al., 1992; Takagi and Shibata, 2000; Miyamoto et al., 2000;

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Takeda et al., 2000). Recently, Tazawa (2004) proposed a revised tectonic framework for the Japanese Islands and divided the pre-Tertiary geological units into the following five terranes; 1) South-Kitakami accretionary terrane (Early Ordovician – Late Devonian with younger cover sequences), 2) Akiyoshi accretionary terrane (Middle – Late Permian), 3) Mino accretionary terrane (Early Jurassic – Early Cretaceous with Cretaceous metamorphism of Ryoke metamorphic belt), 4) Shimanto accretionary terrane (Late Cretaceous – Neogene), 5) Hida-Abukuma nappe terrane (Late Jurassic to Cretaceous nappes including Paleozoic – Mesozoic metamorphic complexes). According to this classification, the Mizukoshi Formation, Mifune Group and Ryuhozan metamorphic complex in the Higo terrane are considered to belong to the South-Kitakami accretionary terrane. The remaining Higo and Manotani metamorphic complexes and the Higo plutonic complex as described below in detail are distinguished as the members of

Fig. 1. Simplified geological map of the Higo terrane.

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the Hida– Abukuma nappe terrane as well as metamorphic complexes of Abukuma and Hida – Oki region (Tazawa, 2004). Remarkable Late Permian to Triassic metamorphism and related plutonic activities are recorded in the Higo and Hida – Oki metamorphic complexes. These events were synchronous with continent-collision metamorphism at the boundaries between the North and South China, and South China and Indochina cratons in East Asia (e.g., Isozaki and Maruyama, 1991; Li et al., 1993; Zhang et al., 1996; Maruyama et al., 1997; Osanai et al., 2001a, 2004; Oh et al., 2004). This contribution briefly reports the geology and the metamorphic process of the Higo and Manotani metamorphic complexes and discusses the possible location of these metamorphic complexes in East Asia before forming a nappe structure. In the following sections mineral abbreviations are mainly after Kretz (1983). 2. Geological outline of metamorphic complexes in the Higo terrane Metamorphic rocks in the Higo terrane are classified into three complexes named the Manotani, Higo and Ryuhozan metamorphic complexes (Karakida, 1992). The Manotani metamorphic complex consists mainly of low-grade greenschist-facies rocks, which preserve earlier high-pressure/lowtemperature (HP/LT) relict minerals, such as alkali-amphiboles and lawsonite (Karakida et al., 1989). The Ryuhozan metamorphic complex is characterized by intense mylonitization and low-grade metamorphism (Yamamoto, 1962), where characteristic chloritoid-bearing pelitic schist is also found (Sakashima et al., 1995). In contrast the Higo metamorphic complex consists of very high-grade metamorphic rocks up to ultrahigh-temperature (UHT) granulite-facies with remarkable anatexite (Yamamoto, 1983; Obata et al., 1994; Osanai et al., 1996, 1998). The boundary between the Manotani and the Higo metamorphic complexes is a thrust fault accompanied by serpentinite intrusion, while that between the Higo and the

Ryuhozan metamorphic complexes is not clear because of the intrusion of the Late Triassic Miyanohara tonalite and the Early Cretaceous Shiraishino granodiorite (Kamei et al., 2000) which are part of the Higo plutonic complex. A simplified geological map of the Higo terrane is shown in Fig. 1 and a schematic geological cross-section of the Higo metamorphic complex and other related metamorphic and plutonic complexes is given in Fig. 2. The Higo metamorphic complex is an imbricate crustal section in which a sequence of metamorphic rocks shows a correlation of increasing metamorphic grade with lowering of structural level from north to south. The complex is composed of various kinds of medium- to high-grade metamorphic rocks with a general metamorphic foliation striking E – W to ENE – WSW and homoclinally steep (50 – 80-) dips (Fig. 1). Pelitic and semipelitic metamorphic rocks (biotite– muscovite gneiss, biotite gneiss, garnet – biotite gneiss, garnet –cordierite – biotite gneiss, garnet –orthopyroxene –biotite gneiss) predominate throughout the sequence with minor amounts of mafic (amphibolite, garnet amphibolite, orthopyroxene– hornblende granulite, orthopyroxene– clinopyroxene – hornblende granulite), intermediate (hornblende – biotite gneiss) and calc – silicate rocks (garnet – clinopyroxene marble, wollastonite-bearing marble). These metamorphic rocks are overlain by pure-limestone nappe (e.g., Mt. Kamakura nappe: Okamoto et al., 1989) (Figs. 1 and 2). The basal part of the metamorphic sequence (southern part) represents an original depth of 23– 24 km where orthopyroxenebearing or garnet – cordierite-bearing S-type tonalites (anatexite) and hornblende or hornblende – clinopyroxene gabbros are distributed as layer-parallel intrusions in high-grade metamorphic rocks. In the northwestern part of the Higo metamorphic complex ultramafic rocks (serpentinite and metamorphosed peridotite) are characteristically observed along shear zones oblique to the metamorphic foliation (Figs. 1 and 2). Some tectonic blocks of highly aluminous high-grade metamorphic rocks (sapphirine – corundum, garnet –corundum, corundum – spinel –cordierite and garnet –spinel granulites) are included in this type of ultramafic rock (Osanai et al., 1998, 2001a).

Fig. 2. Schematic geological cross-section of the Higo metamorphic terrane. Metamorphic zoning of Manotani and Higo metamorphic rocks (A to E) is after Obata et al. (1994). Every metamorphic rocks and older intrusions of gabbro, ultramafic rocks and Miyanohara tonalite are intruded by the younger Shiraishino granodiorite.

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The Manotani metamorphic complex consists mainly of mafic schists derived from MORB-type basalt with minor amounts of pelitic/psammitic schist, red chert, dolerite, gabbro and serpentinite. Metamorphic structure generally indicates E – W foliations and steep N dips with metamorphic grade increasing from north to south. The dismembered ophiolitic original rock constitution of the Manotani metamorphic complex is clearly different from the protoliths of the Higo metamorphic complex as described above and these two complexes are bounded by thrust fault with intrusion of serpentinite. However, metamorphic conditions of the basal part (southern part) of the Manotani metamorphic complex grade into those in the upper portion (northern part) of the Higo metamorphic complex (Obata et al., 1994; Maishima and Osanai, 1995). Therefore the main metamorphism of these two complexes was simultaneous and took place after tectonic juxtaposition. 3. Overview of metamorphism The metamorphic rocks of the Higo and Manotani metamorphic complexes that are exposed on the surface can be subdivided into five metamorphic zones (zone A to zone E) from top to base (i.e., from north to south) on the basis of mineral parageneses of pelitic rocks. Greenschistfacies mineral assemblages in zone A (Karakida et al., 1989) give way to amphibolite-facies assemblages in zones B and C, which in turn are replaced by granulite-facies assemblages in zones D and E. The regional distribution of each zone is given in Obata et al. (1994) and Maki et al. (2004). The metamorphosed peridotites in zones C and D, which occur along the shear zone oblique to the regional metamorphic foliations, include remarkable sapphirine- and corundumbearing highly aluminous granulites of restitic nature (Osanai et al., 1998). No exposure of these aluminous granulites as layers or masses is observed on the surface except as blocks in metamorphosed peridotite. P – T conditions of these aluminous granulites indicate higher-temperature and -pressure conditions as the high-grade extension of the metamorphic field gradient made through zones C, D and E. Therefore we consider that the high-aluminous granulites are situated beneath zone E and some rocks from this highgrade portion were trapped in peridotite intrusions during regional shearing, and we define the high-aluminous granulites as zone F situated at deeper portion of the crustal level than zone E (Osanai et al., 2001a). A schematic cross-section including zone F is shown in Fig. 2. Zone A corresponds to the Manotani metamorphic complex and zones B to F accord to the Higo metamorphic complex. Progressive changes in metamorphic grade from zone A to zone F are explained through the following reactions; Zone Zone Zone Zone Zone

A to B: Phe + Chl = Ms + Bt + Qtz + Vapor, B to C: Ms + Qtz = Sil + Kfs + Vapor, C to D: Bt + Sil + Qtz = Grt + Crd + Kfs + Vapor, D to E: Bt + Qtz = Opx + Grt + melt (or Kfs + Vapor), E to F: fluorine-rich Phl + Sil + Qtz T Grt T Pl = Spr + Crn + Crd + melt.

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Pelitic schists from zone A contain lawsonite characteristically. Partial melting has occurred in zones D and E, where Bt-free leucocratic melt-pods (in zone D: Fig. 3) and orthopyroxenebearing S-type tonalite magma were generated (in zone E). Changes in mineral parageneses of mafic metamorphic rocks are equivalent to pelitic metamorphic rocks. Zone A of the Manotani complex is characterized by sodium-poor actinolite + epidote + chlorite T hornblende assemblage in which pumpellyite + sodic actinolite + epidote + chlorite paragenesis is also found. Crossite and riebeckite are remaining as relict minerals inside in sodic actinolite core. These crossite, riebeckite, sodic actinolite and pumpellyite are overgrown by sodium-poor actinolite and hornblende (e.g., Karakida et al., 1989). Mafic metamorphic rocks from zone B show hornblende + plagioclase T actinolite T epidote T quartz and those from zones C and D have hornblende + plagioclase T garnet T quartz and hornblende + plagioclaseT clinopyroxene T quartz assemblages. The metamorphic rocks from zone E and zone F (as blocks in meta-peridotite) show orthopyroxene + hornblendeT clinopyroxene T quartz and orthopyroxene + spinel T gedrite assemblages, respectively. Prograde and retrograde metamorphic processes in each zone are investigated by reaction textures, chemical zoning of minerals and mineral inclusion in porphyroblasts. For example, garnet – cordierite – sillimanite – biotite gneiss from zone D contains large garnet porphyroblasts (up to 30 mm in diameter) showing a chemical zoning (Fig. 3). The dark reddish core portion of the garnet includes many kinds of mineral inclusions, such as staurolite, tourmaline, ilmenite, rutile, quartz, apatite and sillimanite (kyanite?) and has low-X Mg (Mg/Fe + Mn + Mg + Ca: ca. 0.12) and high-X Ca (Ca/Fe + Mn + Mg + Ca: ca. 0.10), while pale pinkish rims include only fibrolite and has high-X Mg (ca. 0.24) and low-X Ca (ca. 0.02) (Fig. 4). This evidence indicates that the garnet porphyroblast has grown during temperature-increasing and pressure-decreasing prograde processes through the staurolite-consuming reactions of staurolite + quartz = garnet + sillimanite (kyanite) + vapor or staurolite = garnet + biotite + sillimanite (kyanite). Cordierite – spinel symplectite surrounding the garnet porphyroblasts is a retrograde product through the reaction of garnet + sillimanite (fibrolite) = cordierite + spinel which proceeds during decompression and cooling. During the retrograde metamorphism, reheating and low-pressure partial melting took place to produce inclusion-free cordierite + melt by the regional intrusion of the Shiraishino granodiorite in zones D and E (Fig. 5). The newly formed inclusion-free cordierite produced by the incongruent melting reaction shows different features from other inclusionrich cordierite formed by sub-solidus metamorphic reactions. Mafic and pelitic metamorphic rocks from zone A also have evidenced of their prograde (temperature-increase, pressuredecrease) history such as relict minerals of sodic amphibole (crossite and riebeckite), pumpellyite and lawsonite that still remain in some porphyroblastic minerals (Karakida et al., 1989) as described above. P– T estimations were carried out using geothermobarometers for pelitic metamorphic rocks from zones C, D, E and F. Applicable geothermometers are garnet –biotite (Thompson, 1976; Pigage and Greenwood, 1982; Perchuk and Lavrent’eva,

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Fig. 3. Garnet – cordierite – sillimanite – biotite gneiss from zone D of the Higo metamorphic complex. A: large zoned garnet is surrounded by later cordierite – spinel symplectite. Biotite-free leucocratic anatectic melt produced during the peak metamorphic stage (ca. 250 Ma). White squared zoned garnet was used for microprobe chemical mapping of Fig. 4, B: photomicrograph of sillimanite included garnet rim that surrounded by cordierite – spinel symplectite, C: abundant spinel and sillimanite inclusions in cordierite from garnet – cordierite – sillimanite – biotite. Scale bar for photomicrographs show 1 mm.

1983), garnet – cordierite (Wells, 1976; Thompson, 1976; Perchuk et al., 1985), garnet – orthopyroxene (Harley, 1984; Sen and Bhattacharya, 1984), garnet – staurolite (Koch-Mu¨ller, 1997) and sapphirine –spinel (Owen and Greenough, 1991). Pressure estimations for each zone were made with geobarometers such as garnet – cordierite (Wells, 1976), cordierite – spinel (Harris, 1981), garnet – orthopyroxene – plagioclasequartz (Newton and Perkins, 1982; Perkins and Chipera, 1985), GRAIL (Bohlen et al., 1983), garnetbiotiteplagioclase – quartz (Hoish, 1991), and the petrogenetic grid of the KFMASH system (e.g., Spear and Cheney, 1989). Estimated peak P – T conditions of each zone are 350– 420 MPa, 640 –720 -C for zone C, 480 – 600 MPa, 740 –820 -C for zone D, 600 – 720 MPa, 800– 870 -C for zone E and 780 –900 MPa, 900 – 960 -C for zone F, respectively (Osanai et al., 1996, 1998, 2001a). A connected line of these peak P –T conditions from zone C of amphibolite-facies to zone F of UHT granulite-facies forms a metamorphic field gradient and indicates a progressive

change in metamorphic grade from upper to lower crustal levels during the specific period (Fig. 6). Fig. 6 also shows the estimated P– T path for zone D and assumed prograde and retrograde P –T paths for the other zones, especially for zone C. Maki et al. (2004) found a new occurrence of garnet – cordierite– sillimanite gneiss from zone C defined by Obata et al. (1994), which contains biotite + K-feldspar + plagioclase T garnet T inclusion-free cordierite T sillimanite. And then they considered an areal expansion of zone D to the north and suggested that the prograde P – T path for this type of rock indicates the low-pressure isobaric heating process (Fig. 6). Inclusion-free cordierite is one of the characteristic minerals indicating low-pressure partial melting during retrograde process for zone D. Therefore the newly defined low-pressure prograde path by Maki et al. (2004) could indicate a re-heating process of rocks from zone C. Zone B has a muscovite + biotite + andalusite assemblage in the pelitic rock and a hornblende+ plagioclase T actinolite T epidote

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Fig. 4. Chemical mapping of zoned garnet from zone D. Mg-increasing and Mn- and Ca-decreasing continuously from core to rim, excepting outer most rim.

assemblage in the mafic rock, which indicates P–T conditions around 200–300 MPa and 550 -C (Figs. 6 and 7: Karakida et al., 1989; Obata et al., 1994). An assumed peak P–T condition of zone A, that has sodium-poor actinolite + chlorite + epidote T hornblende in mafic rocks as described above, could also be assumed as ca. 200 MPa and 450–500 -C according to the NCFMASH petrogenetic grid (e.g., Winter, 2001). Therefore, peak metamorphic conditions of zones A and B would be confirmed as the lower extension of a progressive metamorphic field gradient (Figs. 6 and 7). Mineral chemical data are available from the corresponding author on request. 4. Geochronological outline of the Higo metamorphic complex Geochlonology of the Higo metamorphic complex has been discussed during the last four decades. Shibata and Yamamoto (1965) first reported K – Ar biotite ages of 106 T 11 and 10 T 89 Ma from zone D or E. From the 1990’s many workers have been trying to get a peak metamorphic age using multi-isotope

methods of K – Ar, Rb – Sr and Sm –Nd for mineral (internal) and whole rock isochrons. Micro analytical ages by SHRIMP (for Zrn) and CHIME (for monazite and zircon) methods are also established. Results are classified into the following five age clusters (Table 1). 1) Paleoproterozoic ages (ca. 1.8 – 2.2 Ga), which were detected only at zircon cores from garnet –cordierite –biotite gneiss and garnet – cordierite tonalite (anatexite), indicate the age of inherited zircon from precursors (Suzuki et al., 1998; Osanai et al., 1999; Sakashima et al., 2003). 2) Cambrian – Devonian ages of ca. 450 – 550 Ma from zircon cores in garnet – cordierite – biotite gneiss also date inherited grains (Suzuki et al., 1998), and an age of 377 Ma from phlogopite in corundum – spinel– phlogopite granulite (Ueda and Onuki, 1969), which forms a part of high-aluminous sapphirine –corundum granulite, has also reported. 3) An Early Permian age (ca. 280 Ma) was determined by a Sm – Nd internal isochron using a carefully separated garnet core, mineral inclusions (sillimanite, ilmenite, rutile, etc) at the core and whole rock. The garnet core also includes prograde

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Fig. 5. Cordierite-bearing anatectic melt pods in garnet – biotite gneiss of zone D. A: cordierite in leucocratic melt portion has characteristically free from mineral inclusions, which indicates an incongruent partial melting reaction, B: photomicrograph of inclusion-free cordierite in melt pod, C: euhedral cordierite and plagioclase occur in melt pod. Scale bar for photomicrographs show 4 mm.

staurolite relics, which would indicate that the garnet core was produced by the staurolite-breakdown reaction during this period (Hamamoto et al., 1999). 4) Late Permian to Triassic ages of the Higo metamorphic complex (260 –230 Ma) are considered to reflect the main metamorphic stage (M1 stage as described below) associated with the progressive metamorphic field gradient described above. Suzuki and Takagi (2000) suggested that this age could be dating detritus derived from granitic precursors by the textural feature of zircon. They argued that a ca. 250 Ma Zrn core in the psammitic gneiss from zone E surrounded by a younger rim (ca. 120 Ma) might reflect a detrital magmatic zircon grain overgrown by a younger rim during the main metamorphic stage. However part of their investigated samples were collected from zones D and E, where partial melting took place during the main metamorphic stage and produced anatectic melt-pods, segregation veins and S-type tonalite masses (258 Ma by SHRIMP in Table 1). Growth pattern of zircon derived from these anatexites would be

similar to magmatic growth zoning. Hornblende gabbro situated in the basal part of zone E is a possible heat source for the prograde metamorphism and anatexis, and has an age of 258 Ma determined by Sm –Nd internal isochron (Hamamoto et al., 1999). Therefore we assume that the Late Permian to Triassic ages from the Higo metamorphic complex are the most acceptable ages of the main metamorphic event. Late Triassic to Early Jurassic ages (ca. 180 –210 Ma) of the Manotani metamorphic complex (zone A) were detected by the K –Ar method using separated muscovite (Nagakawa et al., 1997), that indicates a retrograde stage passing across 350 -C after the peak metamorphism of zone A (ca. 200 MPa, 450– 500 -C) by the isotopic closure temperature of muscovite for the K – Ar system. 5) Early Cretaceous ages (ca. 100 –120 Ma) of the Higo metamorphic complex are completely comparable to those of the Higo plutonic complex (Table 1) except a Hbl-whole rock Sm – Nd internal isochron age of the Miyanohara

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Fig. 6. Pressure – temperature-time path of the Higo metamorphic complex. Peak metamorphic conditions of each zone, which show progressive metamorphic field gradient at ca. 250 Ma, are estimated by geothermobarometers and a petrogenetic grid (e.g., Holdaway, 1971; Spear and Cheney, 1989; Le Breton and Thompson, 1988) for pelitic metamorphic rocks. A nearly isothermal decompression prograde path, which might be caused by the gabbroic intrusion (underplating) of 258 Ma, and retrograde path are identified especially in zone D. Re-heating and low-pressure partial melting took place during the retrograde process associated with intrusion of the younger Shiraishino granodiorite (122 Ma) in zones D and E. Peridotite intrusions containing aluminous restitic granulite blocks of zone F were active along the shear zones during the peak metamorphic stage, and intrude into zones C and D. Shaded area indicates a low-pressure prograde P – T path of garnet- and cordierite-bearing pelitic gneiss from zone C (Maki et al., 2004), which would be regarded as a re-heating process as well as zone D affected by Shiraishino granodiorite.

Fig. 7. Simplified pressure – temperature-time paths of the Higo and Manotani metamorphic complexes based on the mafic metamorphic rocks. A peak P – T condition of zone A from the Manotani metamorphic complex adds to the progressive metamorphic field gradient determined by pelitic rocks from the Higo metamorphic complex, which forms a complete progressive P – T trend from zone A to zone F during peak metamorphic stage (M1: ca. 250 Ma). Broken arrow shows an assumed metamorphic field gradient of M0 high-pressure metamorphic stage before peak metamorphism (M1). Lined area shows presumed P – T conditions of zone A during M0-stage, where pumpellyite – actinolite – epidote – chlorite, crossite and riebeckite were stabilized in mafic rock and lawsonite in pelitic rock. White and shaded circles show age-unknown and age-known, respectively. Relevant petrogenetic grids are after Bucher and Frey (1998) and Winter (2001). Related P – T path of the north Dabie terrane shows the similar evolutional process (Zhang et al., 1996).

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Table 1 Recent geochronological results from the Higo and Manotani metamorphic complex and the Higo plutonic complex Complex Proterozoic

Higo MC Higo MC Higo MC Cambrian – Devonian Higo MC Higo MC Early Permian Higo MC Late Permian – Triassic Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Manotani MC Manotani MC Higo PC Early Jurassic Manotani MC Manotani MC Manotani MC Manotani MC Cretaceous Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Higo MC Higo PC Higo PC Higo PC Higo PC Higo PC Higo PC Higo PC Higo PC

Age (Ma) Mineral/Isochron

Method

Rock name

References

2155 ca. 1790 ca. 1940 450 – 550 377 279 261 258 258 ca. 250 248 239 235 229 226 214 206 211 185 182 182 176 123 117 116 110 104 103 103 99 122 105 104 103 ca. 110 107 106 100

SHRIMP SHRIMP CHIME CHIME K – Ar Sm – Nd (M) Sm – Nd (M) SHRIMP Sm – Nd (M) CHIME Sm – Nd (W) Sm – Nd (M) Rb – Sr (W) Rb – Sr (W) Sm – Nd (M) K – Ar K – Ar Sm – Nd (M) K – Ar K – Ar K – Ar K – Ar CHIME SHRIMP CHIME K – Ar Rb – Sr (M) Rb – Sr (M) K – Ar Rb – Sr (M) Rb – Sr (W) K – Ar Rb – Sr (M) K – Ar SHRIMP Rb – Sr (M) Rb – Sr (M) K – Ar

Grt – Crd – Bt gneiss Grt – Crd tonalite (anatexite) Grt – Crd – Bt gneiss Grt – Crd – Bt gneiss Crn – Spl – Phl granulite Grt – Crd – Bt gneiss Grt – Crd – Bt gneiss Grt – Crd tonalite (anatexite) Hbl gabbro pelitic/psammitic gneisses? Spr – Crn granulite Grt – Crd – Bt gneiss Amphibolite/Skarn Grt – Opx – Bt tonalite (anatexite) Grt – Crd – Bt gneiss Psammitic schist Psammitic schist Miyanohara tonalite Pelitic schist Pelitic schist Pelitic schist Pelitic schist Spr – Crn granulite Grt – Crn – Bt gneiss Grt – Crn granulite Bt gneiss Spr – Crn granulite Grt – Crd – Bt gneiss Hbl gabbro Amphibolite/Skarn Shiraishino granodiorite Shiraishino granodiorite Shiraishino granodiorite Shiraishino granodiorite Miyanohara tonalite Miyanohara tonalite Miyanohara tonalite Miyanohara tonalite

Sakashima et al. (2003) Osanai et al. (1999) Suzuki et al. (1998) Suzuki et al. (1998) Ueda and Onuki (1969) Hamamoto et al. (1999)) Hamamoto et al. (1999) Osanai et al. (1999) Hamamoto et al. (1999) Suzuki et al. (1998) Osanai et al. (2001a) Hamamoto et al. (1999) Yamaguchi and Minamishin (1986) Osanai et al. (1993) Hamamoto et al. (1999) Nagakawa et al. (1997) Nagakawa et al. (1997) Kamei et al. (2000) Nagakawa et al. (1997) Nagakawa et al. (1997) Okamoto et al. (1989) Nagakawa et al. (1997) Suzuki et al. (1998) Sakashima et al. (2003) Suzuki et al. (1998) Shibata and Yamamoto (1965) Osanai et al. (2001a) Hamamoto et al. (1999) Hamamoto et al. (1999) Yamaguchi and Minamishin (1986) Kamei et al. (2000) Kamei et al. (2000) Nakajima et al. (1995) Nakajima et al. (1995) Sakashima et al. (2003) Kamei et al. (2000) Nakajima et al. (1995) Nakajima et al. (1995)

Zrn (C) Zrn (C) Zrn (C) Zrn (C) Phl Grt (C) – Incl – WR Grt (C) – Incl Zrn (R) Hbl – FF – WR Zrn (C) WR Grt (R) – FF – WR WR WR Grt (R) – Incl Ms Ms Hbl – WR Ms Ms Ms Ms Mnz Zrn (R) Mnz Bt Bt – FF – WR Bt – FF – WR Hbl Bt – WR WR Hbl Bt – WR Bt Zrn (R) Hbl – WR Bt – WR Bt

Abbreviations: MC, metamorphic complex; PC, plutonic complex; (C), core; (R), rim; Incl, inclusion minerals; FF, felsic fraction; WR, whole rock; (M), mineral isochron; (W), whole rock isochron.

tonalite (211 Ma in Table 1; Kamei et al., 2000). The Shiraishino granodiorite intruded into the Higo metamorphic complex and the Miyanohara tonalite at around 120 Ma (Kamei et al., 2000), where regional contact metamorphism and re-heating took place at part of the Higo metamorphic complex up to ca. 750 -C (Fig. 2). This is also evidenced by the occurrence of cordierite-bearing melt produced by low-pressure partial melting in zones D and E (Fig. 5). Both the metamorphic and plutonic rocks underwent retrograde cooling after the Shiraishino granodiorite intrusion of the Higo metamorphic complex, and they passed ca. 300 -C (isotopic closure temperature of biotite for K – Ar and Rb – Sr systems) simultaneously. The upper crustal parts of zones A and B would have passed these low-closure temperatures earlier than the lower crustal zones C – F during up-thrusting to form the regional nappe structure suggested by Tazawa (2004), for example zone A passed through this temperature at around 210– 170 Ma and zone D at ca. 100 Ma as above.

5. Discussion and concluding remarks 5.1. Tectonometamorphic evolution of the Higo metamorphic complex The peak P– T conditions of each metamorphic zone in the Higo and Manotani metamorphic complexes show a representative andalusite – sillimanite type of low-pressure progressive field gradient from greenschist-facies in zone A (A-1 in Fig. 7) up to UHT granulite-facies in zone F (F-1) through amphibolite- (zone C: C-1) and low-pressure granulite-facies (zone D: D-1 and zone E: E-1). This is considered to be of Late Permian to Early Triassic age (260 – 230 Ma), which we designate as the M1 metamorphic stage (Fig. 7). The prograde metamorphic history before reaching the peak M1-stage are identified only in zones A and D. Mineral constituents of the peak condition in zone A represent a low-pressure greenschist-facies assemblage, in which pumpellyite + sodic actinolite and lawsonite are still remaining as relict minerals indicating the higher-pressure

Y. Osanai et al. / Gondwana Research 9 (2006) 152 – 166

161

Table 2 Tectonometamorphic evolution of the Higo metamorphic complex Deformation stage

Deformation and structure

D0

D1

D2

D3

D4

Layer – parallel dextral shearing Asymmetrical intrafolial fold, sheath fold and S – C structure with mineral lineation E – W trending folding Open to closed microscopic to mesoscopic fold, macroscopic antiform and synform Seismic and aseismic faulting E – W trending fault with cataclasite, foliated cataclasite and pseudotachylite Thrusting Marble nappe with cataclasite, foliated cataclasite and fault gouge

Tectonics

Metamorphism

Tectonic thickning of supracrustal precursors Collision of micro continents in East Asia South China and North China cratons?

Initial LTHP metamorphism? Juxtaposition of the Higo and Manotani complexes Early prograde metamorphism Peak metamorphism with forming a progressive field P – T gradient Partial melting and formation of S-type tonalites Early retrograde metamorphism

Top-to-the-east displacement

N – S compression

Retrograde metamorphism

Top-to-the-south displacement N – S compression Nappe tectonics Top-to-the-south displacement N – S compression Nappe tectonics Uplifting

Retrograde metamorphism passing across ca. 350 -C (zone A) Retrograde metamorphism

Uplifting

D5

Sinistral shearing NNE – SSW trending fault

Transverse faulting

conditions of the previous metamorphic stage (A-0: M0-stage). Inclusions of crossite and riebeckite inside in sodic actinolite could also be stable under the P– T conditions around A0 or higher-pressure conditions of blueschist-facies (Fig. 7). In zone D staurolite-breakdown and garnet-forming reaction is identified during prograde metamorphism (D-0V) of ca. 280 Ma (Osanai et al., 1996), but no sign of an earlier high-pressure metamorphic stage (D-0 at M0-stage?) is discovered yet. The newly found ilmenite –clinopyroxene symplectite in garnet – amphibolite (Nishiyama et al., 2004) and garnet-bearing clinopyroxenite (Maki and Nishiyama, 2004) may indicate high- and ultrahigh-pressure (UHP) conditions, which would be equivalent to M0 metamorphic stage. The retrograde processes are also preserved in rocks from zones A and D as mentioned above.

Re-heating and local partial melting with formation of Crd-melt Retrograde metamorphism passing across ca. 300 -C (zone D) Retrograde metamorphism

Igneous activity

Age

¨ca. 280 Ma

Gabbro and ultramafic rocks

ca. 260 – 230 Ma

Miyanohara tonalite (I-type) without any relation to Higo metamorphic rocks

ca. 210 Ma

210 – 170 Ma

Shiraishino granodiorite

ca. 120 Ma

ca. 106 – 100 Ma

Endo et al. (1997) summarized the tectonic evolution of the Higo metamorphic complex and they recognized six-stages of deformation (D0 to D5). The D0-deformation stage is equivalent to the prograde and peak metamorphic stages (M0 – M1: before 230 –260 Ma). D1- to D5-stages indicate the tectonic movements during the retrograde metamorphism before the Shiraishino granodiorite intrusion (ca. 120 Ma); D1stage shows top-to-the-east displacement and then N – S compression and top-to-the-south displacement with the formation of a regional nappe structure among D2- and D4stages. Final sinistral shearing and NNE-SSW transverse faulting are recognized as a D5-stage. Continuous thrusting and uplifting also took place between D4- and D5-stages. A brief summary of the tectonometamorphic processes in the Higo metamorphic complex are shown in Table 2.

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5.2. The Higo metamorphic complex in correlation with other East Asian metamorphic complexes The main metamorphism of the Higo and the Manotani metamorphic complexes associated with a low-pressure type progressive P – T field gradient would have taken place during Late Permian and Triassic times (260 – 230 Ma) and their P– T evolution paths including prograde and retrograde processes show clockwise trajectories. Late Permian– Triassic high-grade metamorphism in East Asia is commonly associated with the collision of the North China– South China and South China – Indochina cratons (e.g., Zhang et al., 1996; Osanai et al., 2004). The original rock constitutions, metamorphic processes, precursor and metamorphic ages, chemical and geochronological characters of related intrusive granitic rocks, and other geological background of the Higo metamorphic complex are exceedingly similar to those of the ‘‘north Dabie’’ high-grade terrane (Zhang et al., 1996), which is different from the UHP Dabie terrane (Table 3). The north Dabie terrane is considered

to have formed by the collision of the North and South China cratons (e.g., Maruyama et al., 1994; Zhang et al., 1996). To consider the origin of the Higo metamorphic complex of Permo – Triassic age before regional nappe formation and opening the Japan (East) Sea, we make comparison with other Permo – Triassic metamorphic and related plutonic rocks in the Far-Eastern Eurasian continent, which are mainly distributed along the cratonic collision boundaries (Fig. 8). The Imjingang Belt (Ree et al., 1996) and Sulu and Dabie Terranes (e.g., Li et al., 1993; Ames et al., 1996; Rowley et al., 1997; Chen et al., 2001) are located between the North and South China Cratons, while the Truong Son Belt (Lepvrier et al., 2004) and Kontum Massif (e.g., Tran Ngoc Nam, 1998; Osanai et al., 2001b, 2004; Carter et al., 2001; Nagy et al., 2001; Nakano et al., 2003) are distributed along the boundary between the South China and Indochina Cratons, the Pha Son Complex (Ahrendt et al., 1993; Singharajwarapan and Berry, 2000) and BentongRaub Suture (Krahenbuhl, 1991; Kwan et al., 1992) are found between the Indochina and Sibumas Cratons, and finally the

Table 3 Compalison of features among the Higo metamorphic complex, Sefuri metamorphic rocks and north Dabie terrane

Metamorphic rock assembly

Matamorphic facies Metamorphic evolution Protolith

Iintusive rocks

Age related to metamorphism

Higo metamorphic complex (Osanai et al., 1996, 1998, 2001a)

Sefuri metamorphic rocks (Owada et al., 2000, 2004)

North Dabie terrane (Zhang et al., 1996)

Hbl – Bt felsic gneiss Grt – Bt gneiss Grt – Crd – Bt gneiss Cpx – amphibolite Grt – Opx – Cpx mafic granulite Grt – Opx – Crd granulite Grt – Cpx calc – silicate granulite Grt – Crn – Spl granulite Spr – Crn – Spl granulite ultramafic rocks (meta-peridotite, serpentinite) migmatite impure marble amphibolite – granulite (partly UHT) clockwise granitic rocks pelitic rocks impure limestone mafic volcanic rocks (MORB) ultramafic rock Mg – Al restite (Spr-bearing) (Syn-metamorphic) peridotite (harzburgite, dunite) Cpx – Hbl gabbro (258 Ma) Grt – Opx tonalite (anatexite) (258 – 229 Ma) Grt – Crd tonalite (anatexite) (239 Ma) (Post-metamorphic) Hbl tonalite (211 Ma) granodiorite (122 Ma) Grt leucogranite (100 Ma) Inherited age 2155 – 1790 Ma Pre-peak metamorphic age 550 – 279 Ma Peak metamorphic age 260 – 230 Ma Retrograde metamorphic age 123 – 99 Ma

Bt felsic gneiss Grt – Crd – Bt gneiss Cpx – amphibolite Opx – Cpx mafic granulite Grt – Opx – Crd granulite ultramafic rock impure marble

Bt – Hbl felsic gneiss Cpx – amphibolite Grt – Opx – Cpx mafic granulite Grt – Opx – Bt felsic granulite Cpx – Opx – Mag – Qtz granulite Grt – Opx – Crd – Os granulite* Grt – Cpx calc – silicate granulite ultramafic granulite migmatite marble

amphibolite – granulite clockwise? granitic rocks pelitic rocks impure limestone mafic volcanic rocks (MORB) ultramafic rock

amphibolite – granulite partly UHT) clockwise TTG complex pelitic rocks impure limestone mafic volcanic rocks ultramafic rock Mg – Al restite (Os-bearing)* (Syn-metamorphic) peridotite (harzburgite, dunite) Grt pyroxenite gabbronorite

*: osumilite in Grt – Opx – Crd granulite (Osanai, unpublished data).

(Syn-metamorphic)

(Post-metamorphic) granodiorite (116 Ma) Grt leucogranite (100 Ma)

(Post-metamorphic) diorite (211 Ma) granite (133 – 122 Ma)

Inherited age

Inherited age 1861 Ma Pre-peak metamorphic age 435 Ma Peak metamorphic age 244 Ma, 224 Ma Retrograde metamorphic age 130 Ma

Pre-peak metamorphic age Peak metamorphic age Retrograde metamorphic age ca. 100 Ma?

Y. Osanai et al. / Gondwana Research 9 (2006) 152 – 166

163

Fig. 8. Distribution of Permo – Triassic metamorphic and related plutonic rocks in East Asia. Subdivision of micro continents (cratons) is after Metcalfe (1999).

Jinshajiang-Aliaoshan Suture (Wang et al., 2000) is identified at East Himalayan Syntaxis zone (Fig. 8). Most of the metamorphic complexes in these collision terranes contain not only the HT/LP high-grade rocks but also HP/LT and UHP rocks. Some small masses of Permo –Triassic metamorphic rocks are also found inside the North China Craton (e.g., Fangshan metamorphic rocks of NW-Beijing) (Ishiwatari, 1999), but they are not associated with any high- and ultrahigh-pressure metamorphic rocks. Therefore the Higo metamorphic complex, with possible evidence of high- and ultrahigh-pressure metamorphism (Maki and Nishiyama,

2004), could be a collision type of metamorphic terrane, probably related geographically to the North – South China boundary. Metamorphic and related plutonic correlations between the Higo metamorphic complex and the Ryoke metamorphic belt form the current basis of the pre-Tertiary tectonic framework in west Japan. It is accepted as being beyond doubt that the Ryoke metamorphic rocks were derived from the Jurassic accretionary complex and the related Cretaceous granitic intrusions have high-Sr initial ratios (SrI: > 0.707). However, the supracrustal precursors of the Higo metamorphic complex were dated as at

Fig. 9. Permo – Triassic tectonic framework of Eastern Asia with relation to the distribution of the high-grade metamorphic rocks (modified from Yin and Nie, 1993).

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least older than Early Permian (ca. 280 Ma) and the granitic rocks in the complex show remarkably low-SrI (< 0.706) and older ages than those from the Ryoke metamorphic belt (Osanai et al., 2000; Kagami et al., 2000). This suggests that the lower crustal compositions to produce the granitic magma could be different (Kagami et al., 2000), which may support the idea that the Higo metamorphic complex could not be the western extension of the Ryoke metamorphic belt. Recently small masses of low-pressure type high-grade metamorphic rocks (garnet– orthopyroxene– cordierite granulite, orthopyroxene – clinopyroxene mafic granulite, etc.) have been found around the mountain ridge of the Sefuri Mountains in northern Kyushu (Owada et al., 2000, 2004). Geochronological data for them are not yet available, but P – T conditions and their metamorphic evolution are similar to those of zones D and E of the Higo metamorphic complex, and related Cretaceous granitic activities (SrI: <0.706) also show a strong correlation to the Shiraishino granodiorite (Owada et al., 2000; Kamei et al., 2000). Owada et al. (2000) suggested that granulites from the Sefuri Mountains could be a transported fragment associated with top-to-the-south displacement and nappe formation in the Higo metamorphic complex. It is concluded preferably that the Higo metamorphic complex including the Manotani metamorphic complex and related high-grade metamorphic rocks located at the Sefuri Mountains in northern Kyushu have strong relationships to the Permo – Triassic metamorphic terranes in East Asia. These metamorphic complexes were displaced from the continental collision metamorphic terrane in East Asia with the formation of a regional nappe during the Late Triassic to Early Cretaceous period, together with the Hida – OKi terrane (Komatsu et al., 1985) (Fig. 9). These concluding remarks also suggest a strong possibility that the Higo metamorphic complex might contain the diamond- and coesite-bearing UHP metamorphic rocks typical of the Dabie and Sulu terranes. Acknowledgements We would like to sincerely thank N. Ono, O. Maishima, D. Doyama, Y. Yoshihara and T. Ando for their assistance during the fieldwork. Thanks are also due to I. Fitzsimons, S. Dasgupta and T. Shimura for critical reading of this manuscript. This work was partly supported by a Grant-in-Aid for Scientific Research from the Ministry of Education, Culture, Sports, Science and Technology, Japan (No. 14340150 and No. 17253005: Y. Osanai). This is a contribution to IGCP-368 and -440. References Ahrendt, H., Chonglakmani, C., Hancen, B.T., Helmcke, D., 1993. Geochronological cross section through north Thailand. J. Southeast Asian Earth Sci. 8, 207 – 214. Ames, L., Zhou, G., Xiong, B., 1996. Geochronology and geochemistry of ultrahigh-pressure metamorphism with implications for collision of SinoKorea and Yangtze cratons, central China. Tectonics 15, 472 – 489. Bohlen, S.R., Wall, V.J., Boettcher, A.L., 1983. Geobarometry in granulites. In: Saxena, S.K. (Ed.), Kinetics and Equilibrium in Mineral Reactions. Springer-Verlag, New York, pp. 141 – 171.

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Further reading Ames, L., Tilton, G.R., Zhou, G.Z., 1993. Timing of collision of Sino-Korean and Yangtze cratons. Geology 21, 339 – 342. Arakawa, Y., Saito, Y., Amakaw, H., 2000. Crustal development of the Hida Belt, Japan: evidence from Nd – Sr isotopic and chemical characteristics of igneous and metamorphic rocks. Tectonopys 328, 183 – 204. Barr, S.M., MacDonald, A.S., 1987. Nan River suture zone, north Thailand. Geology 15, 907 – 910. Cho, D.L., Suzuki, K., Adachi, M., Chwae, U., 1996. A preliminary CHIME age determination of monazite from metamorphic and granitic rocks in the Gyeonggi massif, Korea. J. Earth Planet. Sci. 43, 49 – 65. Cobbing, E.J., Mallick, D.I.J., Pitfied, P.E.J., Teoh, L.H., 1986. The granites of southeast Asian tin belt. J. Geol. Soc. (Lond.) 143, 537 – 550. Faure, M., Lin, W., Monie, P., Le Breton, N., Scha¨rer, U., 2001. Tectonics of Dabieshan-Sulu (E. China) structural constraints on the exhumation of ultrahigh-pressure continental rocks. Sixth Int. Eclogite Conf., (abst.), pp. 36 – 38. Lepvrier, C., Maluski, H., Van Vuong, Nguyen, Roques, D., Axente, V., Rangin, C., 1997. Indochina NW-trending sear zones within the Truong Son belt (Vietnam): 40Ar 39Ar Triassic ages and Cretaceous to Cenozoic overprints. Tectonophys 283, 105 – 127. Liew, T.C., Page, R.W., 1985. U – Pb Zircon dating of granitoids plutons from the West Coast Province of Peninsular Malaysia. J. Geol. Soc. (Lond.) 142, 515 – 526. Liu, Y., Xu, S., Li, S., Jahn, B.M., Jiangm, L., Wu, W., Che, G., 2001. Petrologic geochemistry, isotopic chronology and cooling history of eclogite from Northern Dabie Mountains, central China. Sixth Int. Eclogite Con., (Abst.), pp. 82 – 83. Oh, C.W., Choi, S.G., Zhai, M., Guo, J., Song, S., Jin, Y.Y., Seo, J.E., 2003. The first evidence of high-P/T metamorphism in Korea and its implication to north east Asia tectonics. Fifth Hutton Symp., (abst.), p. 112. Tran Ngoc Nam, Sano, Y., Terada, K., Toriumi, M., Van Quynh, Phan, Le Tien Dung, 2001. First SHRIMP U – Pb zircon dating of granulites from the Kontum massif (Vietnam) and tectonothermal implications. J. Asian Earth Sci. 19, 77 – 84.