In the Mariana forearc, horst and graben structures are well developed in the outer forearc basement, which is composed of both island arc and oceanic crust-mantle rocks. A zone of dome-shaped diapiric seamounts, which are composed mainly of serpentinized peridotites, formed on the basement in the outer forearc regions. Serpentine minerals in peridotites from both diapiric seamounts and basement are mostly chrysotile and/or lizardite. Antigorite, however, is rarely found in peridotites recovered from Conical, Big Blue, Celestial, and South Chamorro Seamounts. Antigorite-bearing peridotites always contain secondary iron-rich olivine and metamorphic clinopyroxene, and antigorite seems to coexist stably with them. Iron-rich secondary olivine (Fo86–90) occurs as overgrowth on the rim or along the cleavage traces of primary olivine (Fo90–92). The assemblage shows high-temperature conditions of serpentinization at ~450–550 °C, whereas chrysotile- and/or lizardite-bearing assemblages occur at ~200–300 °C. In antigorite-bearing samples, chrysotile and/or lizardite veins both predating and postdating antigorite formation are recognized. This may reflect a complex process of tectonic cycling of shallow mantle wedge serpentinized peridotites to depth and then back again to the surface.
Serpentinized peridotites are often found along the trench-landward slope of the western Pacific margin. Fisher and Engel (1969) first reported serpentinized peridotites from the landward slope of Tonga Trench, and after the second half of the 1970s, it was shown clearly that serpentinized peridotites and related ophiolitic rocks are widely exposed on the trench- landward slope of the Izu-Bonin and Mariana areas (e.g., Honza and Kagami, 1977; IGCP Working Group, 1977 [IGCP—International Geological Correlation Programme]). On the basis of topographic analysis, Fryer et al. (1985) made an important contribution in finding a zone of dome-shaped serpentinite seamounts formed on the basement where horst and graben structures dominate in the Mariana forearc regions. Ocean Drilling Program (ODP) Leg 125 drilled two serpentinite seamounts, Conical Seamount in the Mariana forearc and Torishima Forearc Seamount in the Izu-Bonin forearc, and obtained the following important results. An analysis of the fluids seeping from the chimneys at the summit of Conical Seamount suggests that the fluids were derived from the dehydration processes of a descending oceanic slab (Fryer et al., 1990; Mottl, 1992). The blueschist facies rock clasts found in the cores recovered at the southern flank of Conical Seamount suggest a blueschist facies metamorphism beneath the forearc (Maekawa et al., 1993). Ishii et al. (1992) studied serpentinized peridotites recovered from the above two serpentinite seamounts, and concluded that the peridotites were residues of extensive partial melting (30%), the last episode of which occurred in the mantle wedge, probably associated with the generation of incipient island arc magma, including boninite and/or arc tholeiite sources. Parkinson and Pearce (1998), however, examined the geochemical nature of these peridotites, and indicated that the forearc is underlain by two types of mantle lithosphere, one being trapped, or accreted oceanic lithosphere, the other being lithosphere formed by subduction-related melting.
The summit of South Chamorro Seamount was drilled during ODP Leg 195 (Shipboard Scientific Party, 2002). D'Antonio and Kristensen (2004) demonstrated that serpentine + brucite paragenesis characterizes the serpentinization of peridotites recovered from Hole 1200A at South Chamorro Seamount, and estimated the upper temperature limit for serpentinization as 200–300 °C. Sand-sized fragments of blue amphibole–rich schists and tremolitechlorite schists were found from the recovered cores at South Chamorro Seamount during ODP Leg 195 (Shipboard Scientific Party, 2002). In Maekawa et al. (2004), these fragments were regarded as metasomatic products between wedge mantle serpentinite and pelagic sediments on subducting plate along the subduction boundary. The analysis of inner structures of serpentinite seamounts using multichannel seismic and bathymetric data suggested that each of the serpentinite seamounts formed by episodic eruptions of mudflows from a central edifice (Oakley et al., 2007).
Recent seismological research has given evidence of widespread serpentinization of the mantle wedge. Based on the seismic reflection-refraction study, Takahashi et al. (1998, 2007) showed that the upper mantle velocity beneath the forearc gradually decreases toward the trench axis to become indiscernible from the velocity of the lower crust in the Izu-Bonin and Mariana subduction systems. Unusually low velocity (7.1 km/s) of the upper mantle beneath the east side of the forearc suggests that a large amount of water is carried down and released by subduction for serpentinization of the mantle peridotites. They demonstrated that the root of the serpentinite diapir on the inner trench wall is a low-velocity mantle wedge that was probably caused by a large amount of water released from the subducting Pacific plate at depths shallower than 30 km. Shimamoto (1985) and Shimamoto et al. (1993) discussed the seismicity and deformation mechanisms in subduction zones and divided a subduction plate boundary into three zones: shallow, intermediate, and deep interfaces. He ascribed the aseismic and decoupled natures of the shallow interface to the existence of enormous amounts of water. The shallow interface may correspond to the low-velocity mantle wedge. The low-velocity wedge at the western side of the trench may indicate the path of a serpentinite diapir (Takahashi et al., 1998). The low electric resistance at the boundary between the arc and subducting plate is due to the high water content of the region, which corresponds to the low-velocity mantle wedge (Toh, 1993).
In this paper we describe the petrologic characteristics of serpentinized peridotites obtained from Mariana forearc, mainly during the R/V Kairei KR06–15 cruise (2006) of the Japan Agency for Marine-Earth Science and Technology (JAMSTEC), and partly during ODP Leg 125 (1989), ODP Leg 195 (2001), and the R/V Yokosuka YK03–07 cruise (2003) of JAM-STEC, and discuss the tectonic significance of these serpentinized peridotites. Although most of the studies on Mariana forearc peridotites deal with relatively small areas, this study deals with peridotites from 11 sites, which cover almost the entire area of the southern Mariana forearc. We intend to explain comprehensively the regional characteristics and significance of serpentinization of Mariana forearc peridotites.
GEOLOGIC OUTLINE AND SAMPLE LOCATIONS
In the Mariana forearc, horst and graben structures developed under an extensional stress field are well developed in the outer forearc, the western limit of which is 60–90 km from the trench. The outer forearc basements consist of both island arc and oceanic crust-mantle rocks, such as serpentinized peridotites, gabbros, volcanics, and their metamorphic equivalents of low-pressure and high-temperature type, and pelagic chert (Meijer et al., 1982; Wood et al., 1982; Bloomer and Hawkins, 1983; Johnson et al., 1991). Diapiric seamounts are scattered on the outer forearc basements. The seamounts are commonly smooth sided, dome shaped, as much as 30 km in diameter with as much as 2 km of relief, and consist mainly of serpentinized peridotites, with subordinate amounts of volcanic, hypabyssal, and plutonic rocks of both island arc and oceanic origins and their high-pressure and low-pressure metamorphic equivalents (Maekawa et al., 1992, 1993). Figure 1 shows the locations where we obtained rock samples. During the KR06–15 Cruise in 2006, seven sites were investigated by the KAIKO 7000II dive (Maekawa et al., 2007a); five sites (369, 370, 371, 374, and 375) were on the flank near summits of dome-shaped diapiric seamounts, and two (372 and 373) were at the fault scarps developed along the edges of horst-like mounds. In addition to serpentinized peridotite samples from these sites, we used the samples obtained from five dives by Sinkai 6500 during YK03–07 Cruise in 2003 (dives 778–780 at the summit of South Chamorro Seamount, 782 at the summit of Celestial Seamount, and 786 at the northeastern steep cliff of Coni-Pack Triangle), and from ODP Legs 125 and 195.
We identified samples collected from eight dome-shaped seamounts and three basement fault scarps for this study. There is a clear difference in constituent rock types between both of the occurrences. The seamounts are rich in serpentinized peridotite; based on the number of recovered rocks, the content is >60%. The fault scarp rock samples, however, contain significant amounts of dolerite, basalt, and gabbro, and contain <40% serpentinized peridotites (Fig. 1).
PETROLOGICAL AND MINERALOGICAL CHARACTERISTICS
Most of the peridotites in the study area were moderately to highly serpentinized, and weathered by seawater to appear brown to brownish red to the naked eye. Some, however, have well-preserved primary mineralogy. Primary olivine, clinopyroxene, and orthopyroxene have locally survived in the rock samples, whereas primary chromian spinel is found in most rocks. Modal analyses of peridotites indicate that harzburgite is predominant and dunite is subordinate. Primary spinel chemistry, however, suggests a possible lherzolite protolith for some highly serpentinized peridotites in North Chamorro and Turquoise Seamounts. Constituent minerals of representative samples of peridotite are given in Table 1.
Minerals were analyzed with a JEOL JSM-840A scanning electron microscope equipped with an Oxford energy-dispersive analytical system (Link ISIS series L200I-S) at Osaka Prefecture University. Accelerating voltage and beam current were kept at 15.0 kV and 0.5 nA, respectively. Corrections were made using the ZAF method (Sweatman and Long, 1969). About spinels, the trivalent iron was estimated from the stoichiometry. Mineral species identification for the serpentine mineral polymorphs was done by polarizing microscope and confirmed by an imaging-plate X-ray microdiffractometer (IP-XRD) with a two-dimensional area detector using graphite-monochromatized Cu-Kα radiation, operating at 40 kV and 40 mA (IP-XRD, RINT RAPID, Rigaku Co.) at the Department of Physical Science, Osaka Prefecture University. Collimators of 0.1 and 0.3 mm diameters were used for analysis of a selected area in thin section by CCD (charge-coupled device) camera. Using the IP-XRD, serpentine polymorphs were easily identified by the diagnostic peaks at 2.525Å (2θ = 35.6°) for antigorite, 2.501Å (2θ = 35.9°) for lizardite, and 2.519Å (2θ = 36.6°) for chrysotile (Fig. 2) (Kohyama, 2007). A Rigaku thermogravimetry-differential thermal analysis (TG-DTA) analyzer at Osaka Application Laboratory, Rigaku Co. Ltd., was also used for part of the samples to identify the three serpentine polymorphs by the differences of dehydration temperatures, that is, in the case of 10 °C/hour increasing temperature ratio, they are ~600 °C for lizardite, 700 °C for chrysotile, and 750 °C for antigorite.
Olivine commonly occurs as a porphyroclastic grain of 1–4 mm length in serpentinized peridotites in the Mariana serpentinite seamounts. It is highly deformed and often has wavy extinction and kink band. An aggregate of fine-grained olivine of <30 μm is also found around porhyroclastic olivine and orthopyroxene. It is thought to have been formed during mylonitization that promoted serpentinization. Both types are chemically equivalent (Fo89–93), and we regard these olivines as primary ones.
Cleavable olivine, which has perfect cleavages on the (100), (010), and (001) crystal planes (Hawkes, 1946), were reported from peridotites of Conical Seamount by Ishii et al. (1992) and Girardeau and Lagabrielle (1992). We found cleavable olivine in peridotites from Big Blue, Celestial, and South Chamorro Sea-mount (Figs. 3A, 3B, and 3E). Iron-rich olivine is found in cleavable olivine-bearing samples from Conical, Big Blue, and South Chamorro Seamounts. It occurs irregularly around the rim and cleavage trace of primary olivine (Figs. 4A–4C), and seems to have formed along the once-existing conduits of fluid. The iron-rich stripes emanate from cracks filled with antigorite in primary olivine (Murata et al., 2009), indicating that the formation of iron-rich olivine is closely related to serpentinization (antigorite formation). These modes of occurrence suggest a possibility that iron-rich olivines formed secondarily during serpentinization. It usually is accompanied with only small amounts of magnetite and characteristically has an iron-rich composition (Fo86–88) (Figs. 5A, 5B). Some of this type of olivine, however, is locally associated with abundant magnetite, and has chemical compositions slightly richer in forsterite content (Fo88–90) (Figs. 4G and 5C). It is noteworthy that the iron-rich olivine is only found in the peridotites with cleavable olivine and is always accompanied with antigorite and acicular clinopyroxene (Figs. 3C, 3D, 3F, 4A, 4B, and 4H).
The representative chemical compositions of olivine are given in Table 2.
Serpentinized peridotites from the Mariana forearc almost always contain chromian spinel, which is partly replaced by magnetite along the rim and cleavage trace during serpentinization. Euhedral spinel is common in the dunite samples, suggesting that they may have crystallized from a melt or formed by melt-mantle interaction. Spinels in peridotites from the Mariana forearc have chemical zonings with Cr- and Fe-rich core and Al- and Mg-rich rim, but those with Cr-rich rim are rarely found. Dick and Bullen (1984) pointed out that the increasing Cr# [100 Cr/(Cr + Al)] of spinel reflects increasing degrees of partial melting in the mantle, and divided peridotites and associated volcanics into three groups on the basis of Cr# of spinels; Type I and Type III contain spinels with Cr# < 60 and Cr# > 60, respectively, and Type II is a transitional group and contains spinels spanning the full range of spinel compositions in Type I and Type III peridotites. The compositions of spinel core in peridotites from the Mariana forearc are plotted in the Mg# [100 Mg/(Mg + Fe2+)]-Cr# diagrams (Fig. 6). In Figure 6, spinels in peridotites from the Mariana forearc are plotted in the range of Type II of Dick and Bullen (1984), and seem to have been formed in the island arc that developed on oceanic crust. Spinels in peridotites recovered from the fault scarps developed in the basement plot in the region with high Cr# ranging from 58 to 91, and those recovered from the dome-shaped seamounts have a wider range of Cr#, from 26 to 75. Peridotites from Turquoise and North Chamorro Seamounts contain aluminous spinels, suggesting the existence of lherzolite, although these rocks are severely serpentinized. In Figure 6, spinels in the peridotites from the fault scarps of each site seem to tend to have a narrow range of Cr#, whereas those from the dome-shaped seamounts of each site plot in the slightly wider range of Cr#. The wide range of Cr# in the seamounts relative to the fault scarps may be explained by the fact that the dome-shaped seamounts had been formed by diapiric rise of serpentinized peridotites with various degrees of melting that were incorporated from the various depths and throughout the pathway of diapiric rise in the subduction zone.
Orthopyroxene occurs as porphyroclasts 1–5 mm in diameter scattered in olivine matrix. Some orthopyroxenes have fabrics of plastic deformation, such as kink bands, bending, and wavy extinction. Orthopyroxene crystals are sometimes surrounded by aggregates of fine-grained olivine neoblasts. They are more or less replaced by chrysotile and/or lizardite, but orthopyroxene pseudomorphs are usually well preserved as bastite texture. Orthopyroxenes are enstatite with compositions of En90–93. The Al2O3 and CaO contents of orthopyroxene vary from 0.59 to 2.16 wt% and from 0.13 to 1.56 wt%, respectively. The Cr2O3 content ranges from 0.06 to 0.84. Representative chemical compositions of orthopyroxene are given in Table 3.
Primary clinopyroxene occurs as a euhedral to subhedral crystal near orthopyroxene. Thin exsolution lamellae of clinopyroxene are commonly found within orthopyroxene crystals. In most cases, clinopyroxene is partly or fully altered to serpentine and/or dusty clay minerals. XMg [Mg/(Mg + Fe2+)] of clinopyroxene ranges from 0.87 to 0.97. The Ca/(Ca + Mg + Fe2+) ratios range from 0.46 to 0.50. The Al2O3 and Cr2O3 contents of clinopyroxene range from 0.56 to 2.19 wt% and from 0.11 to 1.24 wt%, respectively.
Secondary clinopyroxene is found as a fine-grained acicular crystal in the peridotites from Conical, South Chamorro, and Big Blue Sea-mounts. It is always associated with antigorite and secondary olivine (Figs. 3F, 4H, and 4I). Secondary clinopyroxenes have lesser Al2O3 contents and slightly higher CaO contents than primary ones (Fig. 7; Table 3). The XMg ratios of secondary clinopyroxene range from 0.94 to 0.98, slightly higher than those of primary one. Both types of clinopyroxene are classified into diopside.
Colorless to pale green amphiboles are found in samples from Conical, Big Blue, Pacman, Twin Peaks, and South Chamorro Seamounts. They commonly occur as euhedral to subhedral prismatic crystals <0.3 mm in length. The sample Leg 125–779A-19R-2, 105–108 cm, contains olivines with fine-grained amphibole (tremolite) inclusions. Most of them have <5 wt% Al2O3, and correspond to tremolite to magnesiohornblende of Leake et al. (1997) (Table 4). The sample Leg 195–1200A-6R-1, 70–76 cm, from South Chamorro Seamount, contains amphiboles with 10–11 wt% Al2O3, which are classified as edenite or pargasite (after Leake et al., 1997). The modes of occurrence suggest that these amphiboles occur as primary phases formed in hydrous mantle, as suggested by Ohara and Ishii (1998). These amphiboles had lost Al components in varying degrees, probably during serpentinization.
Serpentine minerals in peridotites from the Mariana forearc are commonly chrysotile and lizardite, which occur as platy and fibrous crystals that replaced olivine, orthopyroxene, and clinopyroxene, and as vein fillings. Hourglass and mesh textures after olivine, and bastite textures after both orthopyroxene and clinopyroxene are commonly found (Figs. 3G, 3H, and 4F). Antigorite was found in peridotites from Conical, Celestial, Big Blue, and South Chamorro Seamounts (Table 1). Antigorite-bearing samples from Conical, Big Blue, and South Chamorro Seamounts characteristically have cleavable olivine and secondary iron-rich olivine. Feather-like crystals of antigorite often cut olivine in random directions (Figs. 3A–3D). The mode of occurrence of antigorite is different from that of chrysotile and/or lizardite in that the latter is always associated with magnetite, but the former is accompanied with only small amounts of magnetite (Fig. 4D–4F). Antigorite often coexists with acicular clinopyroxene (Fig. 4I). In antigorite-bearing samples, chrysotile and/or lizardite also occur in a vein or matrix as early-stage and later stage serpentine minerals. Chrysotile and/or lizardite veins cut by antigorite veins are often recognized, and antigorite veins cut by chrysotile and/or lizardite veins are also common (Figs. 4D, 4E). Antigorite has high XMg ratios and tends to be slightly rich in Si in comparison with chrysotile and/or lizardite (Fig. 8; Table 4).
Colorless brucite is found in veins of an antigorite-bearing peridotite from Conical Seamounts (779A-22R-2, 9–14 cm). It is always accompanied with magnetite. We consider that this brucite did not coexist stably with antigorite, because it occurs only as a vein mineral. Pale brown iron-rich brucite described by D'Antonio and Kristensen (2004) from South Chamorro Seamount is also found as a vein filling in peridotites from Conical and South Chamorro Seamounts. This brucite does not chemically coincide with that of ideal brucite, with the Mg(OH)2 chemical formula, and contains significant amounts of Fe(OH)2: FeO contents are 9.83–29.5 wt%. Pale brown brucite does not coexist with antigorite, but is probably in equilibrium with chrysotile and/or lizardite. Representative chemical compositions of brucite are given in Table 5.
Fine-grained Mg-rich phlogopite (XMg = 0.95–0.96) was reported from peridotite of Conical Seamount (Sample 779A-11R-1, 94–96 cm) by Parkinson and Pearce (1998). In this study, phlogopite was newly found in peridotite from South Chamorro Seamount (Sample 195–1200B-1W-1, 92–100 cm) as 1–5 μm inclusions of primary olivine and clinopyroxene (Table 5). The phlogopites often have smaller K2O contents than the idealized formula due to later stage serpentinization or weathering, but those from less altered ones contain significant amounts of K2O, as high as 8.05 wt%. The XMg values of phlogopites in the peridotite from South Chamorro Seamount (0.88–0.95) are lower than those from Conical seamount reported from Parkinson and Pearce (1998).
Origin of Iron-Rich Olivine
Antigorite-bearing peridotites in the Mariana forearc commonly contain primary olivine with well-developed cleavages (cleavable olivine). Cleavages of primary olivine are always filled with antigorite films. The secondary iron-rich olivine occurs irregularly along the rim or cleavage trace of cleavable olivine (Fig. 4A–4C), and as aggregates of subhedral small grains with magnetite (Fig. 4G). Iron-rich olivine seems to have formed secondarily along the irregular bands in olivine that are thought to have been developed conduits of fluid. Where antigorite fills cracks in primary olivine crystal, iron-rich stripes emanate from the cracks along cleavages in the olivine (Murata et al., 2009). Although antigorite has a Mg-rich composition (XMg = 0.95–0.96) in comparison with primary olivine (XMg = 0.90–0.92), in most cases antigorite crystals replaced after primary olivine are accompanied with only small amounts of magnetite. These lines of evidence suggest that a part of the excess iron component resulted from antigorite formation after primary olivine may have dissolved in an H2O-rich fluid that was responsible for the crystallization of the secondary iron-rich olivine. Magnetites are locally abundant along the grain boundaries of fine-grained secondary olivines (Fig. 4G) that have relatively Mg-rich compositions, and are plotted in the area between primary olivine and secondary olivine that is almost free from magnetite association (Fig. 5C). Frost (1985) indicated that serpentinization of peridotite yields an intense reduction state because the formation of magnetite from olivine continuously consumes oxygen from fluid. Iron-rich olivine may have been formed in association with magnetite under the reduction state. Iron-rich olivine is only found in antigorite-bearing peridotite, and always occurs in intimate association with antigorite. Although some iron-rich olivines seem to have been cut by feather-like antigorite, we consider that the iron-rich olivine and antigorite coexist stably at least part of the periods during serpentinization.
Cleavable olivine often develops in the aureoles thermally metamorphosed by intrusion. Uda (1984), in his study on cleavable olivine developed in the thermally metamorphosed Oeyama peridotites, Japan, found that cleavage represents antigorite films that are continuous to dislocation cell walls, and concluded that the dislocation cell structure formed through the annealing recovery process, and that the cell walls were successively invaded by the antigorite film, resulting in the formation of the cleavage. In the cleavable olivines found in the Mariana forearc peridotites, cleavages are always filled with antigorite film, suggesting a strong correlation between the origin of cleavage and antigorite recrystallization. Ohara and Ishii (1998) also reported cleavable olivine from peridotites containing talc, tremolite, and antigorite as secondary phases from the landward slope near the southern end of Mariana Trench, and indicated the similarity to the Oeyama peridotites in that cleavable olivine-bearing peridotites in both areas contain secondary hydrous phases. There is a possibility that cleavable olivine is ubiquitous in antigorite-bearing peridotites. As any heat source is difficult to consider in the Mariana subduction system, cleavable olivine in the Mariana forearc may have formed under a high-pressure environment, as suggested by Kuroda and Shimoda (1967). Further study is necessary to solve the origin of cleavable olivine in the Mariana forearc.
Chrysotile and lizardite are ubiquitous serpentine minerals in peridotites throughout the Mariana forearc, and were found from all the investigated sites. Mineral associations of chrysotile and/or lizardite ± brucite of serpentinized peridotites from the Mariana forearc indicate that the temperatures of serpentinization are ~200–300 °C (D'Antonio and Kristensen, 2004) (Fig. 9). Antigorite-bearing rocks are rare, but were found from Conical, Big Blue, Celestial, and South Chamorro Seamounts. The possibility that once-prevailing antigorite in serpentinized peridotites was completely obliterated to chrysotile and/or lizardite can probably be refuted, because most of the chrysotile- and/or lizardite-bearing peridotites retain mesh and bastite textures filled with chrysotile and/or lizardite. In most cases, these textures are wholly obliterated by feather-like antigorite crystals in antigorite-bearing peridotites.
Antigorite is stable at higher temperatures than chrysotile (Iishi and Saito, 1973; Evans et al., 1976), and lizardite is considered to have the same P-T stability field as chrysotile (Peacock, 1987). Antigorite commonly coexists with acicular clinopyroxene (diopside) and secondary olivine, suggesting high-temperature serpentinization at ~450–550 °C (Fig. 9). In antigorite-bearing samples, we recognized both early and later stage chrysotile veins; that is, the early stage chrysotile was cut by antigorite veins and the later stage chrysotile cut antigorite veins (Figs. 4D, 4E). This suggests that the chrysotile and/or lizardite formations occurred before and after antigorite formation. The early stage formation of chrysotile and/or lizardite is considered to have occurred before the subduction event or initial stage of progressive serpentinization within the subduction zone. However, the later stage chrysotile and/or lizardite indicate that the later low-temperature serpentinization probably took place during the uplift stage. In the Mariana forearc, serpentinite diapirs are considered to have generated along the subduction boundary, where serpentinization took place due to the water supplied from the hydrated sediments on top of the subducting Pacific plate. As antigorite-bearing assemblages favor the deep high-temperature portion in the subduction zone, the antigorite-stable region is farther from the trench axis along the subduction boundary than the chrysotile- and/or lizardite-stable region. In antigorite-bearing samples, chrysotile and/or lizardite veins both predating and postdating the antigorite formation are recognized. This may reflect a complex process of tectonic cycling of shallow mantle wedge serpentinized peridotites to depth and then back again to the surface.
Blueschist facies rocks were found in metabasite clasts recovered from Conical Seamount ODP Leg 125 (Maekawa et al., 1993) and from South Chamorro Seamount during ODP Leg 195 (Shipboard Scientific Party, 2002; Maekawa et al., 2004). The stable association of antigorite + diopside + olivine is well known in blueschist and eclogite facies terrains (e.g., Scambelluri et al., 1991). The finding of the same assemblage in the Mariana forearc peridotites shows that the stable existence of the assemblage actually formed in a subduction zone environment. We cannot explain the reason why we did not find the antigorite + brucite assemblage that is located at an intermediate region between the chrysotile- and/or lizardite-stable and antigorite + olivine–stable regions in Figure 9. As only small amounts of peridotite samples have been examined in this study, more comprehensive study in the Mariana forearc is necessary to understand the more precise nature of serpentinization in subduction zones.
We are indebted to the captain and crew of the R/V Kairei KR06-15 cruise and the Kaiko 7000II operation team for their help. We are grateful to M. Scambelluri and an anonymous reviewer for critical reading and valuable comments on the manuscript. We thank K. Ozawa and Y.O. Mohammad for helpful and fruitful discussions. K. Murata is grateful to H. Goto and S. Hada for their heartfelt and continuous encouragements. This research was supported by a Grant-in-Aid for Scientific Research from Japan Society for the Promotion of Science to K. Murata (16730431).