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Journal of Petrology | Volume 45 | Number 2 | Pages 369-389 | 2004
© Oxford University Press 2004; all rights reserved
Petrology of Peridotite Xenoliths from Iraya Volcano, Philippines, and its Implication for Dynamic Mantle-Wedge Processes
1 DEPARTMENT OF EARTH SCIENCES, FACULTY OF SCIENCE, KANAZAWA UNIVERSITY, KANAZAWA 920-1192, JAPAN
2 INSTITUTE OF GEOSCIENCES, FACULTY OF SCIENCE, SHIZUOKA UNIVERSITY, SHIZUOKA 422-8529, JAPAN
RECEIVED NOVEMBER 12, 2002; ACCEPTED AUGUST 14, 2003
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Peridotite xenoliths entrained in calc-alkaline andesites from the Iraya volcano, Philippines, were petrologically examined to constrain the nature of the mantle-wedge materials and processes. They can be classified into two types: C-type (coarse-grained type) and F-type (fine-grained type) peridotites. C-type peridotites are mostly coarse-grained (olivine,
1 mm across) harzburgites with porphyroclastic to protogranular textures but include subordinate dunites. F-type peridotites are fine-grained (olivine, 60–70 µm across). Secondary orthopyroxenes that replace olivine and sometimes show radial (spherulitic) aggregation are very common in F-type peridotites and, subordinately, in C-type peridotites, in which the total amount of orthopyroxene increased in volume. Fine-grained olivine in F-type peridotites characteristically has minute glass and chromian spinel inclusions. Mineral chemistry is clearly different between the two types of peridotite: olivine is around Fo91–92 and Fo89–91 in C-type and F-type peridotites, respectively. The Cr/(Cr + Al) atomic ratio (Cr number) and Fe3+/(Cr + Al + Fe3+) atomic ratio of chromian spinel are 0·2–0·3 and <0·1, respectively, in C-type peridotites, and 0·4–0·7 and around 0·1, respectively, in F-type peridotites. The secondary orthopyroxenes are appreciably lower in Al2O3, Cr2O3 and CaO than the primary ones. A textural transition from C-type to F-type peridotites can be observed; coarse olivine becomes recrystallized into fine grains through subgrains that preserve the previous coarse texture. The C-type harzburgites are similar in mineral chemistry to arc-type harzburgites, e.g. mantle xenoliths from the Japanese island arcs, and may represent samples of the sub-arc lithospheric mantle. The C-type harzburgites beneath the Iraya volcano may have been strained and deformed during oblique subduction of the South China Basin. A silicate melt rich in SiO2, H2O and Fe, possibly derived by fractional crystallization from a primitive arc magma, assisted the recrystallization of the C-type peridotites to the F-type peridotites with metasomatic chemical modification. Oblique subduction is common in arc–trench systems, suggesting that F-type peridotite formation may be common within the mantle wedge. KEY WORDS: mantle wedge; peridotite; metasomatism; Iraya volcano; Philippines
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Samples of the sub-arc mantle, represented by peridotite xenoliths entrained in arc magmas, are rare relative to mantle samples from non-arc settings, i.e. from oceanic hotspots and continental rift zones (e.g. Nixon, 1987
). This means that there is a paucity of xenolith-based direct petrological information about the mantle wedge relative to other tectonic settings. Hence the rare examples of arc-derived peridotite xenoliths need to be investigated systematically and in detail to explore the nature of mantle-wedge materials and processes.
Peridotite xenoliths of possible mantle-wedge origin have been described from the Japanese island arcs (e.g. Takahashi, 1978; Aoki, 1987; Abe, 1997; Abe et al., 1998; Arai et al., 1998, 2000), the Colorado Plateau (e.g. Smith & Riter, 1997; Smith et al., 1999), the Cascades, USA (Brandon & Draper, 1996; Ertan & Leeman, 1996), Mexico (Luhr & Aranda-Gómez, 1997), Papua New Guinea (Grégoire et al., 2001; McInnes et al., 2001; Franz et al., 2002) and Kamchatka (Kepenzhiskas et al., 1995; Arai et al., 2003). In this study we focus on the peridotite xenoliths hosted in arc-type andesite of the Iraya volcano, in the Luzon arc (Richard, 1986; Maury et al., 1992). Among the Iraya peridotite xenoliths extremely fine-grained peridotites [F-type of Arai & Kida (2000)] predominate over coarse-grained types (C-type). Peridotite xenoliths with similar characteristics are also known from the Avacha volcano, Kamchatka, and it has been proposed that the fine-grained peridotites are characteristic of the mantle wedge beneath island arcs (Arai et al., 2003). Their distinctive characteristics have not been observed in other tectonic settings (e.g. oceanic hotspots and continental rift zones), but are probably common to mantle-wedge peridotites. In a previous paper (Arai & Kida, 2000) we presented basic petrographical and mineral chemical data and referred to a possible deserpentinization (= dehydration recrystallization from serpentinite) origin for the F-type peridotites. Here we present a new interpretation, based on a more detailed petrological study of the peridotite xenoliths from the Iraya volcano, and discuss the petrological characteristics of the mantle wedge. We focus especially on the origin of the F-type peridotites, based on petrological and fabric analyses in the context of the tectonic situation of the mantle wedge.
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GEOLOGICAL AND TECTONIC BACKGROUND |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Batan is the main island of the Batanes Province, bounding the northernmost territory of the Philippines (Fig. 1). The volcanoes of Batan belong to the Babuyan Segment, the least evolved of four segments of the Luzon arc (Defant et al., 1989
, 1990). Batan is located at the junction of the western and eastern chains of the Taiwan–Luzon arc (Yang et al., 1996). The Babuyan Segment has evolved on the western part of the Philippine Sea plate, which is subducted by the South China Sea plate and the Eurasian plate along the Manila Trench (e.g. Lallemand et al., 2001) (Fig. 1). The underthrust plate has a very high angle of subduction or is even overturned beneath Batan (Yang et al., 1996; Lallemand et al., 2001). The Philippine Sea plate is moving northwestward with a velocity of about 7 cm/year relative to the Eurasian plate (Seno, 1977).
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Batan comprises three volcanoes, Mahatao, Matarem and Iraya (Fig. 1) with different ages (Richard et al., 1986a
, 1986b). Mahatao volcano is the oldest; its eruption started during the Late Miocene in the central part of Batan. Matarem volcano in the southern part of the island has been strongly dissected and its volcanic products are covered by volcanics from the two younger volcanoes, sediments and coral reefs. Matarem volcano is Pliocene to Early Pleistocene in age. Iraya volcano has been active since the Late Pleistocene in the northern part of the island (Fig. 1). The volcanic rocks from Batan are andesitic to basaltic (Richard et al., 1986a). Gabbroic cumulate xenoliths and hornblende megacrysts have been reported from pyroclastics of the Mahatao volcano (Richard et al., 1986a).
Xenoliths are especially abundant in recent pyroclastics of calc-alkaline series lavas (1480 years BP) erupted from Iraya volcano (Richard et al., 1986a, 1986b), collected mainly from cliffs at Song-Song Bay and Balugan Bay. They are ultramafic (peridotitic) to mafic (gabbroic) in composition, rounded to subangular in shape and are up to 25 cm across. Subordinate peridotite xenoliths have also been found in volcanics from Matarem volcano. Xenoliths of basement crystalline schists are common in the lavas from Matarem volcano but are very rarely found in the lavas from Iraya volcano. Fine-grained (F-type) peridotite xenoliths are predominant over coarse-grained (C-type) types (Arai et al., 1996; Arai & Kida, 2000). Typical C-type peridotite xenoliths, with porphyroclastic to protogranular texture, are very rare, consisting of about 4% of all the xenolith samples examined.
The volcanics hosting the peridotite xenoliths were analyzed by XRF at Kanazawa University. They contain 49–60 wt % SiO2 and are mostly andesites with relatively high K2O contents, belonging to the high-K series. They plot in the calc-alkaline field and around the boundary between calc-alkaline and tholeiitic series on a SiO2–FeO*(total FeO)/MgO diagram.
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PETROGRAPHY OF THE XENOLITHS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Arai et al. (1996)
classified the peridotite xenoliths from Iraya volcano into two types in terms of grain size, C-type (coarse-grained type) and F-type (fine-grained type). The two types of peridotite are very different in appearance, petrography and mineral chemistry (Arai et al., 1996). C-type and F-type peridotites have a light olive green color and very pale yellowish green color, respectively, in hand specimen. Some xenoliths are intermediate between the two types: coarse-grained peridotite is either cut or enclosed by a fine-grained part. This suggests a transformation from C-type to F-type peridotites as described in detail below.
Modal proportions of minerals were determined by point-counting, involving 2000–3000 points covering the whole area of a thin section (Fig. 2). Some uncertainty is expected for the C-type peridotite xenoliths because of their small sample size. F-type peridotites are often too fine-grained for point-counting analysis; consequently, only F-type peridotites with relatively coarse-grained textures were analyzed by the point-counting method (Arai & Kida, 2000) (Fig. 2).
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The peridotite xenoliths of both types from Iraya have a hornblendite selvage of which the thickness is highly variable from sample to sample (Arai et al., 1996
). They are occasionally entirely enclosed in hornblende gabbro, although the hornblende gabbro is never in direct contact with the peridotite. In extreme cases, hornblendite encloses small angular peridotite clasts up to 1 cm across, which are partly disintegrated and darker-colored (dull yellowish green), indicating partial digestion and chemical modification.
C-type (coarse-grained) peridotites
C-type peridotites are mostly harzburgites (Figs 2 and 3), and usually exhibit protogranular to weakly porphyroclastic textures (Fig. 3a and c). The rare dunites have a tabular equigranular texture (Fig. 3e and f). The volume ratio of clinopyroxene/pyroxenes is mostly less than 0·1 in C-type harzburgites. C-type harzburgites are occasionally higher in orthopyroxene content than abyssal peridotites (Fig. 2) partly because of the presence of orthopyroxene-rich pockets that are olivine-orthopyroxenite in mode and are <0·5 cm across (e.g. sample 124-1 of Fig. 2). The C-type harzburgites rarely contain clinopyroxene-rich bands. Olivine in harzburgites is up to 5 mm across, and is clear but partly turbid as a result of glass inclusion trails (Fig. 3b). Olivine and orthopyroxene frequently show wavy extinction or kink bands (Fig. 3a), and orthopyroxene porphyroclasts, up to 1 cm across, commonly contain thin exsolution lamellae of clinopyroxene especially in their central part (Fig. 3d). Clinopyroxene is anhedral, fine-grained and small in amount; it is commonly associated with orthopyroxene porphyroclasts. It is subhedral and is selectively turbid in many samples. Chromian spinel is anhedral and brown-colored in thin section (Fig. 3c). Plagioclase and hydrous minerals are totally absent.
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F-type (fine-grained) peridotites
F-type peridotite xenoliths contain green-colored speckles up to 1 cm across, which can be identified as concentrations of minute grains of chromian spinel in thin section. F-type peridotites are similar in their modal mineralogy to C-type peridotites (Fig. 2); foliation occurs in some samples. Olivine is around 60–70 µm across and contains minute spherical inclusions of orbicular glass with chromian spinel and bubbles (Schiano et al., 1995
) (Fig. 4a–d). Chromian spinel typically occurs in fine-grained aggregates of various shapes and is dark brown to black (Fig. 4e and f), often accompanied by glass. This glass is interstitial to the spinel aggregate and is larger in size than other types of glass. The chromian spinel often exhibits pull-apart textures, suggesting that the original coarse spinel grains were flattened and split into pieces perpendicular to the foliation plane (Fig. 4e and f). Globules of Fe–Ni sulfide are characteristically found in F-type peridotites. Plagioclase is very rarely found as an anhedral grain interstitial to olivine (Arai et al., 1996). Very small amounts of amphibole with a pale greenish color occur in the F-type peridotites.
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The F-type peridotites frequently contain two types of coarser olivine. One is clear and euhedral to anhedral in shape, and is medium in size up to 1 mm across (Fig. 4g and h). It is characteristically free from strain and is equivalent to the ‘tablet olivine’ described in peridotite xenoliths from kimberlites (Boullier & Nicolas, 1975
). The other is as coarse as olivine in the C-type peridotites, and is turbid owing to minute glass inclusions (Fig. 4g). This type of olivine occasionally includes the tablet olivine above (Fig. 4g). This coarse olivine is considered to be a relic of olivine from a C-type peridotite protolith (Fig. 3). Fine olivine grains were annealed to form medium strain-free olivine, or more favorably, the coarse strained olivine grains, with tablet olivine (Fig. 4g), were selectively recystallized into fine grains (Fig. 4h).
Textural variation from C-type to F-type peridotite xenoliths is shown in Fig. 5. In C-type peridotites, relatively coarse grains of olivine show triple junctions with straight grain boundaries. The grain-size distribution of olivine is nearly log-normal and its mean grain size is 800 µm (Fig. 6). In some intermediate-type peridotites, these coarse grains of olivine contain uniform subgrains (Figs 5b and 6). It should be noted that subgrains occur in old coarse grains and they have remarkably uniform sizes from one grain to another (Figs 5b and 6). The size distribution of these subgrains is log-normal and the mean subgrain size is 50 µm (Fig. 6). In F-type peridotites, extremely fine olivine grains occur and their grain sizes are extremely uniform in each xenolith (Figs 5c and 6). The grain-size distribution is log-normal and the mean grain size is 70 µm (Fig. 6).
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Olivine CPO patterns
Olivine CPO (crystallographic preferred orientation) of both C-type and F-type peridotites was measured by scanning electron microscopy (SEM) on highly polished thin sections using a JEOL 5600 system equipped with electron back-scattered diffraction (EBSD). A total of 290 and 299 olivine crystal orientations were determined respectively and the computerized indexation of the diffraction pattern was visually checked for each orientation. Although the structural reference frame is unknown in these samples, the measured olivine CPO is presented on equal area, lower hemisphere projections, where the maximum density of the [100] axis was aligned east–west and the maximum density of the [010] axis north–south (Fig. 7).
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Olivine CPO in the C-type peridotite sample is characterized by strong concentrations of [100] and [010] axes (Fig. 7a). The CPO occurs as a single crystal-like point maximum, which is similar to a typical (010)[100] pattern (e.g. Michibayashi & Mainprice, 2004
, fig. 5). The olivine CPO in the F-type peridotite sample is also characterized by a single crystal-like point maximum similar to that in the C-type peridotite (Fig. 7b). However, the concentrations of axes are significantly weaker that those in the C-type peridotite.
Secondary orthopyroxene
Secondary orthopyroxenes are commonly found in C-type and subordinately in F-type peridotite xenoliths from Iraya. The secondary orthopyroxene has ragged boundaries with olivine and contains irregular-shaped fine-grained olivine grains (Figs 3e, f and 8a, b, d). The fine-grained olivine inclusions have the same crystallographic orientation as the surrounding coarse olivine, suggesting a replacement origin for the orthopyroxene (Fig. 8b and d). Small secondary orthopyroxene grains also form within coarse primary olivine (Fig. 8c). Enrichment in orthopyroxene can be demonstrated in some of the samples. A coarse-grained harzburgite (72-1) contains secondary orthopyroxene replacing olivine but has still preserved the primary minerals and texture (Fig. 3d). It contains about 21% of primary orthopyroxene and about 20% of secondary orthopyroxene, the total orthopyroxene being over 40% by volume (Fig. 2). In this harzburgite, the primary orthopyroxene has also begun to be converted to finer secondary orthopyroxene laths around the rim (opx-2 of Fig. 3d), as described in sub-arc peridotite xenoliths from the Avacha volcano, Kamchatka (Arai et al., 2003). Two relatively coarse-grained samples (60-12 and E19-12) were initially dunites with equant chromian spinel, and have only secondary orthopyroxene replacing olivine (Fig. 3e and f). The primary texture is well preserved, mainly composed of coarse to medium grains of olivine that are only partly replaced by relatively fine orthopyroxene (Fig. 3e and f). The amount of secondary orthopyroxene reaches 16 and 25 vol. %, respectively, for 60-12 and E19-12 (Fig. 2).
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The secondary orthopyroxene is sometimes unusual in texture, showing spherulitic or radial aggregates, no sign of deformation, and no exsolution of clinopyroxene (Fig. 8e). It is clearly distinguished from ordinary mantle orthopyroxene similar to that in the coarse-grained harzburgites (Fig. 3a–d). McInnes et al. (2001)
and Arai et al. (2003) described metasomatic orthopyroxenes with exactly the same texture in sub-arc harzburgite xenoliths from the Lihir volcano, Papua New Guinea, and from the Avacha volcano, Kamchatka, Russia. As noted by Arai et al. (1996) and Arai & Kida (2000), spherulitic orthopyroxene has been found as porphyroblasts in deserpentinized peridotites from thermal aureoles around granitic intrusions (see also Arai, 1974, 1975; Matsumoto et al., 1995). Phlogopite with pale brownish colors is sometimes associated with these secondary orthopyroxenes, especially in F-type peridotites. Clinopyroxene is occasionally associated with the secondary orthopyroxene, especially with that recrystallized from primary orthopyroxene.
Host andesites
The xenoliths were entrained mainly by calc-alkaline andesites with phenocrysts of plagioclase, hornblende, augite, biotite, olivine and occasionally hypersthene. Plagioclase is optically zoned but its form is varied from euhedral to subhedral. Some of the plagioclase is clear and some is turbid with numerous glass inclusions. Hornblende is euhedral to subhedral, and is brown to dull green in thin section. Opacite rims are sometimes observed around hornblende phenocrysts. Augite is euhedral to subhedral and pale greenish in color. Magnetite inclusions are common. Olivine is euhedral to round in shape, and coarse euhedral grains enclose brownish euhedral chromian spinel. Orthopyroxene is relatively fine, if present, and frequently has a reaction rim of clinopyroxene. The groundmass is intersertal, with plagioclase, clinopyroxene, magnetite and glass.
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MINERAL CHEMISTRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Minerals and glasses were analyzed with a JEOL electron microprobe (JXA8800) at the Center for Co-operative Research of Kanazawa University (accelerating voltage 15 kV and beam current 12 nA) and with a JEOL 8800 superprobe at the Tokyo Institute of Technology (accelerating voltage 15 kV and beam current 12 nA). Special caution was taken in the NiO analysis of olivine, using 25 kV accelerating voltage, 20 nA beam current and a longer counting time (100 s instead of 20 s for other elements). Ferrous and ferric iron contents of chromian spinel were calculated assuming spinel stoichiometry. Cr number is Cr/(Cr + Al) atomic ratio of chromian spinel. Mg number is Mg/(Mg + total Fe) atomic ratio for silicates and is Mg/(Mg + Fe2+) atomic ratio for chromian spinel. The minerals of the C-type peridotites are almost homogeneous in chemistry, except in samples strongly affected along grain boundaries by the host magma. The minerals in F-type peridotites sometimes show grain-by-grain chemical heterogeneity, and the minerals are too small and too turbid to examine intra-grain chemical heterogeneity in some samples (e.g. Fig. 4a and e). Only the core compositions of the minerals were considered in this study. Representative analyses are listed in Table 1.
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Olivine is Fo91–93 and the Cr number of spinel ranges from 0·3 to 0·6 in C-type peridotites (Table 1; Fig. 9). Dunites have slightly more magnesian olivine (Fo91–93) than harzburgites (Fo91–92). The Cr number of spinel increases steeply with an increase in the Fo content of olivine within the olivine–spinel mantle array, a spinel peridotite restite trend (Arai, 1994
), in C-type peridotites (Fig. 9). Olivine is Fo89–91 and the Cr number of spinel ranges from 0·4 to 0·8 in the F-type peridotites (Fig. 9). It is noteworthy that F-type peridotites have lower Fo content of olivine on average than C-type peridotites despite the higher Cr number of spinel (Fig. 9). Olivine composition is not appreciably different between the intact part and the metasomatized part in which a large amount of secondary orthopyroxene has formed at the expense of olivine (60-12 and E19-12 of Table 1; Fig. 3e and f). The NiO content of olivine ranges from 0·4 to 0·5 wt % in C-type harzburgites (Table 1). It is, however, variable in F-type peridotites: it is sometimes high, 0·5–0·6 wt % (e.g. sample 52 of Table 1), and is sometimes low, around 0·3 wt % (e.g. sample 1 of Table 1). The Fe3+/(Cr + Al + Fe3+) atomic ratio of spinel is low, around 0·05 in C-type peridotites (Fig. 10), and is slightly higher, around 0·1, in F-type peridotites than in C-type peridotites (around 0·05) (Fig. 10). The Cr–Al–Fe3+ ratio of chromian spinel gradually changes from the C-type to F-type peridotites: the Fe3+/(Cr + Al + Fe3+) ratio of the former spinel tends to increase sharply with increase in Cr number (Fig. 10). This transitional spinel is found in peridotites intermediate between the C-type and F-type peridotites mentioned above. Spinel with the chemical signature of F-type peridotite is frequently found in C-type harzburgites with secondary orthopyroxene (e.g. sample A of Table 1). This spinel compositional trend is very different from that related to chemical modification by the hornblendite selvage (Fig. 10). It is noteworthy that the Mg/(Mg + Fe2+) atomic ratio of spinel is slightly higher at a given Cr number in F-type than in C-type peridotites (Arai & Kida, 2000).
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Clinopyroxene is chromian diopside with more than 0·5 wt % of Cr2O3 in C-type peridotites (Fig. 11). Clinopyroxene is slightly poorer in Al2O3, Cr2O3 and Na2O on average in F-type peridotites than in C-type peridotites (Fig. 11). The Na2O content of clinopyroxene is variable from 0 to 1 wt % in individual grains but is low, <0·6 wt % on average, for each sample of C-type peridotite (Fig. 11). Clinopyroxenes in gabbros and as phenocrysts in the host andesite are clearly distinguished from those in peridotites in having higher TiO2 and lower Cr2O3 contents (Fig. 11). Clinopyroxenes in hornblendite selvages, as well as interstitial cpx in peridotites adjacent to the selvages, are intermediate in composition between the peridotite and gabbro/andesite clinopyroxenes (Fig. 11).
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Secondary orthopyroxenes (opx-2 in Table 1), sometimes exhibiting radial aggregation, are characterized by low contents of CaO, Al2O3 and Cr2O3 relative to primary opx in C-type peridotites (Arai & Kida, 2000
) (Table 1). The secondary orthopyroxene replacing olivine (e.g. opx-2 of Fig. 3f) tends to be lower in Ca, Al and Cr than that recrystallized from primary orthopyroxene (e.g. opx-2 of Fig. 3d). The secondary orthopyroxenes are very similar in chemistry to those in metasomatized peridotite xenoliths from the Avacha volcano, Kamchatka (Arai et al., 2003) and from the Colorado Plateau (Smith & Riter, 1997; Smith et al., 1999). The textural and chemical features (especially the radiating form and low Ca content) of the secondary orthopyroxene are very similar to those of metasomatic orthopyroxene in deserpentinized peridotites (Arai & Kida, 2000). Hornblende and phlogopite are generally low in TiO2, <1·2 wt % and <2·8 wt %, respectively. Rare plagioclase is very calcic and is An94–98.
Glass compositions
Schiano et al. (1995) reported the compositions of glasses mainly included in olivine. Complementary to their data, the glasses interstitial to chromian spinel in F-type peridotite and metasomatized C-type harzburgite were analyzed by electron microprobe (Tokyo Institute of Technology) for major elements (Table 2). They are highly silicic and contain 60–66 wt % of SiO2, 2·4–4·0 wt % of Na2O and <3·8 wt % of K2O, and are similar in major-element chemistry to the glasses that occur mostly as inclusions in olivine (Schiano et al., 1995) except for the high analytical totals of our data. This may be due to a difference in volatile contents depending on the mode of occurrence: glasses completely included by minerals [primary inclusions of (Roedder, 1984)] may have higher volatile contents than those from the secondary inclusions analyzed by Schiano et al. (1995). The presence of H2O and almost complete absence of CO2 and other volatiles in the glasses was preliminarily determined by IR microspectroscopy. Normative quartz content varies from 14 to 26 wt %, and the normative quartz/(normative quartz + hypersthene) weight ratio is high and remarkably constant, ranging from 0·71 to 0·75 (Table 2).
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Thermobarometry
We calculated equilibrium temperature for the peridotites using the two-pyroxene geothermometers of Wells (1977)
and Wood & Banno (1973) for pyroxene pairs adjacent to each other. They yield no systematic difference between the C-type and F-type peridotites, giving the same temperature range, 950–990°C for the Wells (1977) geothermometer. The thermometer of Wood & Banno (1973) gives basically the same range but the nominal temperatures are higher by about 100°C than the Wells' temperatures. Arai et al. (2003) also reported a similar equilibrium temperature, about 900°C based on Wells (1977), for both the C-type and F-type peridotite xenoliths from the Avacha volcano, Kamchatka. This result is apparently contradictory to the generally low contents of Ca and Al in the secondary orthopyroxene (Arai & Kida, 2000; Arai et al., 2003). The secondary orthopyroxene is, however, relatively high in Ca and Al if accompanied by clinopyroxene, thus indicating relatively high equilibrium temperatures. Arai & Kida (2000) reported a slightly higher Mg number of spinel cores and lower 
between olivine cores and spinel cores at a given Cr number of spinel, where 
is the apparent partition coefficient normalized to a constant Fe3+ ratio (0·05) after Evans & Frost (1975). Although not examined in detail, this possibly resulted from heating after entrainment by a magma that formed gabbroic selvages at depth before entrainment of the xenolith by the present host andesite. Because of the smaller grain sizes of both olivine and chromian spinel in the F-type peridotites, Mg and Fe2+ may have diffused sufficiently, changing the core compositions, even after heating for only a short period (see Ozawa, 1983). The coexistence of calcic plagioclase (An94–98) with magnesian olivine in some of the F-type peridotites indicates that some peridotites were derived from shallow mantle depths (<1 GPa) (e.g. Kushiro & Yoder, 1966). |
DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Characterization of C-type peridotites
The C-type peridotites from Iraya are different from abyssal harzburgites (Hébert et al., 1983
; Dick, 1989; Cannat et al., 1990; Arai & Matsukage, 1996; Dick & Natland, 1996) in having characteristic orthopyroxene enrichment (Fig. 2). They are also different from the most common ophiolitic harzburgites, for example mantle harzburgites from the Oman ophiolite, in which orthopyroxene is around 20% in volume (e.g. Lippard et al., 1986; Kadoshima, 2002). Instead they are similar to some sub-arc harzburgites; e.g. harzburgite xenoliths from Kurose and Noyamadake, the SW Japan arc, both in mineral chemistry and mode (Fig. 2; Arai et al., 1998, 2000). The relatively low Na2O content of discrete clinopyroxenes is also one of the characteristics of some sub-arc peridotites and abyssal peridotites (Arai, 1994). In summary, we consider that the C-type peridotites from Iraya reflect the petrographical and mineral chemical signatures of the mantle wedge.
Origin of F-type peridotites
Arai & Kida (2000) concluded that the fine-grained peridotites were formed by fluid metasomatism, or alternatively, but less possibly, by deserpentinization of serpentinite in the mantle wedge. The similarity of the characteristic radial aggregate of orthopyroxenes supports a deserpentinization origin for the F-type peridotites. However, we also re-examined the F-type peridotites to see if they could have been transformed from C-type peridotites, assisted by melt migration, through dynamic recrystallization processes.
Transition from C-type to F-type peridotites
In C-type peridotites, coarse grains of olivine contain uniform subgrains, indicating that the coarse grains have been dynamically recrystallized. The size of the subgrains is remarkably uniform from one grain to another, suggesting that the recrystallization mechanism is subgrain rotation (e.g. Passchier & Trouw, 1996). Furthermore, the grain-size distributions in Fig. 6 show that the mean size of subgrains is slightly smaller than the mean grain size of olivine in the F-type peridotites. This size difference between the subgrains and fine olivine grains in the F-type peridotites has also been documented in experimental studies (e.g. Jung & Karato, 2001). Therefore, this suggests that F-type peridotites could have resulted from deformation of C-type peridotites. The transitional process can be illustrated from observations of appropriate samples, where fine-grained olivine aggregates defined by similar crystallographic orientations preserve previous coarse olivine microstructures (Fig. 5b). This shows that the coarse olivine grains in C-type peridotites in Fig. 5a were dynamically recrystallized into aggregates of far smaller subgrains as a result of a subgrain rotation mechanism as shown in Fig. 5b. Further rotation of subgrains because of increasing strain resulted in their weak crystallographic orientations and finally produced F-type peridotites (Fig. 5c).
The grain-size distributions of both C- and F-type peridotites are log-normal and their microstructures are rather uniform. Therefore, assuming that their grain sizes represent steady-state grain sizes, we can estimate the flow stress by a grain-size paleopiezometer. We use the stress versus recrystallized grain-size relationship of Jung & Karato (2001). The estimated flow stress yields 40 MPa for F-type peridotites, but the mean grain size of olivine in C-type peridotites is too coarse to estimate the flow stress by this paleopiezometer.
Although the structural reference frame is unknown, it is noted that the CPO patterns of both C-type and F-type peridotites show a similar pattern to the {0kl}[100] system, which is the most commonly activated slip system in naturally deformed peridotite (e.g. Nicolas & Poirier, 1976). The overall CPO strengths in the F-type peridotites are remarkably weak compared with those in the C-type peridotites (Fig. 7). This may be predominantly due to the subgrain rotation recrystallization, which tends to weaken strong maxima by rotating the crystals away from the ‘ideal’ positions (e.g. Heidelbach et al., 2003).
Minute inclusions of glass and chromian spinel are very common in olivine, especially within its central part, in the F-type peridotites (Fig. 4a and c). This type of inclusion is categorized as a ‘primary inclusion’ (e.g. Roedder, 1984), suggesting entrapment of melt/fluid during the growth of the host fine olivine. This is in strong contrast to the trail of inclusions cutting the coarse olivine in the C-type peridotites (Fig. 3b). This type of inclusion is ‘secondary’ (Roedder, 1984), and was formed along cracks after the formation of the host olivine. The melt/fluid invaded after the formation of the C-type peridotites and during the formation of the F-type peridotites from C-type protoliths, suggesting a dynamic recrystallization of the C-type peridotites assisted by this melt/fluid. Downes (1990) recognized a preference for mantle metasomatism in deformed or sheared parts of peridotite mantle xenoliths.
Fluid or melt for the metasomatic agent?
The orthopyroxene, which is especially characteristic of the fine-grained peridotites, was most probably formed by reactive replacement of olivine by a Si-rich melt/fluid. The relatively low contents of CaO, Al2O3 and Cr2O3 in the secondary orthopyroxene from the Iraya xenoliths are also characteristic of secondary orthopyroxenes from Avacha (Arai et al., 2003), Lihir (McInnes et al., 2001) and the Colorado Plateau (Smith & Riter, 1997; Smith et al., 1999) that have been interpreted to have formed by metasomatism by aqueous fluids of slab origin. Aqueous fluids in equilibrium with peridotite under high-pressure and high-temperature conditions can be reactive with olivine to form orthopyroxene at lower-pressure conditions (e.g. Nakamura & Kushiro, 1974; Stalder et al., 2001; Mibe et al., 2002). Direct information on the nature of the metasomatic agent involved in the formation of the secondary orthopyroxene is not available. The secondary orthopyroxene itself is not accompanied by glass, although glasses are more frequently found in F-type peridotites where the secondary orthopyroxene is most common.
Relatively low partition coefficients (KD values) of Mg–Fe, from 0·1 to 0·3 (mostly 0·1 to 0·2), between glass and olivine are obtained from pairs of host olivine and glass inclusions (Schiano et al., 1995). Our data also yield low KD values, from 0·1 to 0·2, for the pairs of glass associated with spinel and olivine in F-type peridotites (Table 2). The KD values depend on the Fe2+/Fe3+ ratio of the glass: we assumed that the Fe3+/(total Fe) atomic ratio is 0, 0·1 or 0·2 in our calculations because the redox state is unknown (Table 2). The KD values of Mg–Fe between glass and olivine demonstrate positive and negative correlations with (MgO + total FeO) and (Na2O + K2O) of the glass, respectively (Table 2). Combined with the KD values of Schiano et al. (1995), which are generally low, this is totally consistent with the tendency of KD to change depending on the alkali content (Falloon et al., 1997; Draper & Green, 1999) and the (MgO + total FeO) content (Kushiro & Walter, 1998) of the melt, although our KD values, assuming the Fe3+/(total Fe) ratio of the glass as 0, 0·1 or 0·2, are slightly lower than the experimental data. The melts of Falloon et al. (1997) and Draper & Green (1997) are anhydrous to slightly hydrous low-degree partial melts of peridotite, and are nepheline-normative even when they are silicic (with around 60 wt % of SiO2). Taking all the characteristics of the glasses into account, the metasomatic agent could have been a silicate melt with a high H2O content. The relatively low KD values for olivine and glass in the F-type peridotites from Iraya (0·1–0·2) may be due to the relatively high contents of normative quartz and H2O in the melt.
The H2O content of the glasses seems to be systematically variable in the Iraya peridotite xenoliths. As a result of the almost exclusive presence of H2O as a volatile, the analytical total of the microprobe analyses of the glass is expected to be lower than 100% depending on the H2O content. The possible H2O content decreases from the primary glass inclusions in the olivine (around 5–10 wt %; Schiano et al., 1995) to the glass associated with the chromian spinel (almost anhydrous in this study, Table 2) through the secondary glass inclusions in olivine (around 5–7 wt %) (Schiano et al., 1995). This difference of H2O content is due to the loss of H2O on quenching of the melt to various degrees depending on the degree of interconnectivity of the trapped melt. The melt that initially invaded the peridotite was high in H2O (?>10 wt %), considering the presence of bubbles in the glass inclusions in olivine (Schiano et al., 1995) as well as the possible complete miscibility between SiO2-rich silicate melts and hydrous fluids at upper-mantle conditions (Bureau & Keppler, 1999). The melt that formed the glasses was possibly saturated with olivine, after formation of the secondary orthopyroxene, and was not reactive with olivine. Overgrowth of olivine was possible, but may not have occurred because any compositional halo has not been detected around the glass inclusions in olivine.
Orthopyroxene spherulites (Fig. 8g and h) can be formed by supersaturation of the (Mg,Fe)SiO3 component in the fluid/melt and/or by its supercooling (e.g. Inoue et al., 2000). Supersaturation in (Mg,Fe)SiO3 component was probably achieved by the contact of silica-oversaturated melt with olivine within the mantle peridotite. Similar conditions may have operated in the deserpentinized peridotites within the contact aureoles of the granitic intrusions, mentioned above. Arai (1975) reported a higher silica content of the deserpentinized peridotites in the orthopyroxene zone than in the other zones, suggesting silica enrichment from the granitic magma. The silica-rich fluid that emanated from the granitic magma invaded the highest-temperature zone (orthopyroxene zone) of the dehydrating serpentinite and produced spherulitic orthopyroxene in contact with olivine.
There is no systematic increase in the amount of orthopyroxene from the C-type to the F-type peridotites (Fig. 2), although the replacement of olivine with orthopyroxene can be, at least locally, observed (Figs 3e, f and 8b, d). In particular samples, however, orthopyroxene enrichment is discernible (Fig. 3d–f). We cannot conclude that there has been silica enrichment of the mantle wedge based on the Iyara xenolith suite. The migrating melts appreciably modified the peridotites in chemistry (Fig. 9), suggesting that the melt was relatively Fe-rich. They were possibly residual melts fractionated from partially solidified primitive arc magmas at deeper levels.
Implications for tectonic setting
In this part of the Luzon–Taiwan arc the tip of the mantle wedge and the overlying crust is being displaced by shearing parallel to the trench, as a result of the strain partitioning of oblique subduction (Fitch, 1972; Pinet & Cobbold, 1992; Aurelio, 2000). This will lead to strain within the mantle-wedge peridotites, and any melts or fluids present may facilitate deformation/recrystallization (Figs 12 and 13). The South China Sea Basin started to subduct along the Manila Trench beneath the Philippine Sea plate around Middle Miocene times (Stephan et al., 1986). The obliquity of subduction at this time was very high between Taiwan and Luzon; Seno & Maruyama (1984) proposed a north-northwestward movement of the Philippine Sea plate at this time. After the change to the present northwestward movement of the Philippine Sea plate the obliquity of subduction lessened. Consequently, the continuous shearing caused by oblique subduction that deformed the lithospheric mantle to form the F-type peridotites from the C-type peridotites concurrent with invasion of melt (Fig. 12) may have ceased before the onset of the recent activity of the Iyara volcano. The migrating melt dispersed into the surrounding C-type peridotites through cracks and formed trails of secondary glass inclusions (Fig. 3b). The strain-free tablet olivine (Fig. 4g) formed by local recrystallization of strained olivine during annealing as a result of a decrease in the obliquity of subduction (see Drury & Van Roermund, 1989).
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The processes of deformation and recrystallization deduced from the peridotite xenoliths from Iraya, Philippines, may be common to all supra-subduction zone mantle wedges. Subduction not orthogonal to the trench (i.e. oblique subduction) is common, especially around the Western Pacific, and transcurrent faults possibly related to the oblique subduction are also common (Fitch, 1972
). F-type peridotite xenoliths were first described from the Avacha volcano, which is located on the volcanic front of the Kamchatka arc (Arai et al., 2003). This is also consistent with the expected location of the strike-slip faults caused by oblique subduction, i.e. around the volcanic front (e.g. Fitch, 1972). |
SUMMARY AND CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL AND TECTONIC...PETROGRAPHY OF THE XENOLITHSMINERAL CHEMISTRYDISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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- Peridotite xenoliths entrained within calc-alkaline andesites from the Iraya volcano, Philippines, can be classified into two types, C-type (coarse-grained) and F-type (fine-grained) peridotites. Harzburgites with porphyroclastic to protogranular textures are predominant over dunites in the C-type peridotites. Secondary orthopyroxene replacing olivine and sometimes exhibiting radial (spherulitic) aggregation is very common in the F-type peridotites and, subordinately, in the C-type peridotites. Glasses included within olivine or interstitial to fine-grained spinel aggregates are common in the F-type peridotites.
- Mineral chemistry is distinctly different between the two types of peridotite: olivine is around Fo91–92 and Fo89–91 in the C-type and F-type peridotites, respectively. The Cr number and Fe3+/(Cr + Al + Fe3+) atomic ratio of chromian spinel is 0·2–0·3 and <0·1, respectively, in the C-type peridotites, and 0·4–0·7 and around 0·1, respectively, in the F-type peridotites. The secondary orthopyroxenes are appreciably lower in Al2O3, Cr2O3 and CaO than the primary orthopyroxene.
- C-type peridotites are similar in mineral chemistry to arc-type harzburgites, e.g. the harzburgite xenoliths from the Japan arcs. The textural transition from C-type to F-type peridotites can be observed under the microscope: coarse olivine (C-type peridotite) is recrystallized to fine grains (F-type peridotite) through subgrains that preserve the previous coarse size of the original grains. Glasses, mainly trapped in F-type peridotites, are silicate melts rich in SiO2, H2O and Fe. The melt may have assisted the transformation of the C-type peridotites to the F-type peridotites.
- The formation of F-type peridotites from C-type peridotites was due to shearing of the mantle wedge by oblique subduction. This may be common within supra-subduction zone mantle wedges because oblique subduction is common.
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ACKNOWLEDGEMENTS |
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We are grateful to G. P. Yumul, Jr, the University of the Philippines, for his arrangement of and assistance in our field research. Crystal orientation measurements were made by K. Kanagawa with an SEM–EBSD system at the Department of Earth Sciences, Chiba University. We thank E. Hellebrand and two anonymous reviewers for their critical comments that improved an earlier version of the manuscript. The editorial handling and comments of K. Ozawa are gratefully acknowledged. A. Ninomiya, K. Kadoshima, N. Abe, M.V. Manjoorsa and C. P. David collaborated with us to collect the samples in the field. E. S. Andal kindly provided some literature on Philippine geology. Y. Shimizu and S. Ishimaru helped us in preparing the manuscript.
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FOOTNOTES |
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* Corresponding author. Telephone: 81-(0)76-264-5724. Fax: 81-(0)76-264-5746. E-mail: ultrasa{at}kenroku.kanazawa-u.ac.jp
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