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Journal of Petrology Volume 42 Number 7 Pages 1279-1299 2001
© Oxford University Press 2001
Petrogenesis of Corundum-Bearing Mafic Rock in the Horoman Peridotite Complex, Japan
DEPARTMENT OF EARTH SCIENCES, KANAZAWA UNIVERSITY, KAKUMA, KANAZAWA 920-1192, JAPAN
RESEARCH SCHOOL OF EARTH SCIENCES, THE AUSTRALIAN NATIONAL UNIVERSITY, CANBERRA, A.C.T. 0200, AUSTRALIA
Received February 14, 2000; Revised typescript accepted November 17, 2000
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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A corundum-bearing Type II mafic rock, within the Horoman peridotite, Japan, was petrologically examined in detail to obtain the P–T paths of the mafic rock as well as of the host peridotite. Of all the mafic rocks documented from the Horoman complex, only the corundum-bearing mafic rock has preserved, at least partly, its high-pressure mineralogy; all of the others have been completely recrystallized at low pressures. The Type II mafic rocks were initially formed at <1·0 GPa as cumulates of olivine, plagioclase and clinopyroxene. Corundum was then formed by metamorphism and/or partial melting of the Type II protolith at higher pressures (>1·5 GPa) than the initial condition of formation. Corundum reacted with clinopyroxene during exhumation of the Horoman peridotite down to the plagioclase stability field. The field and petrographical observations of the Type II mafic rocks (± corundum) coupled with published isotopic data suggest a complicated spiral-like P–T trajectory for the Horoman peridotite. The Type II protolith was formed at low pressure within the peridotite at the time of initial formation of the Horoman peridotite as a residue from primitive mantle at
830 Ma. The Type II mafic rocks, as well as the surrounding peridotite, then experienced subduction to the garnet stability field. Finally, the Horoman complex ascended a second time from the garnet peridotite to the plagioclase peridotite stability field. The Horoman peridotite is an example of multiple recycling of peridotite within the mantle. KEY WORDS: Horoman peridotite complex; mafic rock; corundum; P–T path; recycling
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Conspicuous mantle heterogeneity, in terms of mafic layers (or pyroxenite–eclogite) in peridotite, is observed in many orogenic lherzolite massifs, for example, Ronda (Dickey, 1970
; Obata, 1980; Garrido & Bodinier, 1999) and Beni Bousera (Kornprobst, 1969) in the Betic–Rifean orogenic belt, Cabo Ortegal (Girardeau et al., 1989) in northern Spain, Lanzo in the Western Alps (Bodinier, 1988), and Lherz in the Pyrenees (Bodinier et al., 1987). These mafic layers vary in thickness from a few centimeters to a few meters and can be sometimes traced over distances of several hundreds of meters. The mafic layers are sometimes concordant with foliation of the peridotite and frequently exhibit folding and boudinage as a result of plastic deformation. Their origins are still controversial (e.g. Allègre & Turcotte, 1986; Garrido & Bodinier, 1999; Takazawa et al., 1999). A heterogeneous mantle relatively enriched in mafic components, including pyroxenite–eclogite, will yield voluminous and compositionally distinctive magmas upon partial melting because of selective fusion of the mafic components (e.g. Hofmann & White, 1982; Hauri, 1996; Hirschmann & Stolper, 1996; Cordery et al., 1997; Takahashi et al., 1998). The products of selective fusion may react with the more refractory wall-rock leading to refertilized, ‘anomalous’ mantle able to act as a distinctive source in later mantle upwelling (Yaxley & Green, 1998). The origins of mafic layers in peridotite complexes are, therefore, very important not only for the origin of mantle heterogeneity but also for magma genesis.
The Horoman peridotite complex, Hokkaido, northern Japan, is well known to have a conspicuous layered structure consisting of various peridotites with minor amount of mafic layers (e.g. Komatsu & Nochi, 1966; Niida, 1984; Takahashi, 1991; Takazawa et al., 1999). The Horoman peridotites experienced partial melting, melt extraction and metasomatism (Obata & Nagahara, 1987; Arai & Takahashi, 1989; Takahashi et al., 1989; Frey et al., 1991; Takahashi, 1991, 1992; Yoshikawa et al., 1993; Takazawa et al., 2000; Yoshikawa & Nakamura, 2000). These processes are considered to be very important to the formation of the layered structure of the Horoman complex. In contrast, Toramaru (1997), Yoshikawa & Nakamura (2000) and Toramaru et al. (2001) suggested that folding of a restite–cumulate unit by deformation produced the layered structure in the Horoman complex. Thus detailed analysis of the precise P–T history of the Horoman complex is critical to our understanding of these mantle processes. The Horoman plagioclase peridotite recorded nearly adiabatic decompression from the garnet peridotite stability field, and was re-equilibrated within the plagioclase peridotite stability field (Ozawa & Takahashi, 1995; Takazawa et al., 1996; Ozawa, 1997; Yoshikawa & Nakamura, 2000). The cores and rims of clinopyroxene porphyroclasts in the plagioclase peridotite are in chemical equilibrium with garnet and with plagioclase, respectively (Takazawa et al., 1996; Yoshikawa & Nakamura, 2000). Symplectites consisting of fine-grained pyroxenes and spinel (some of which contain plagioclase) have been interpreted as garnet pseudomorphs (Kushiro & Yoder, 1966; Takahashi & Arai, 1989; Ozawa & Takahashi, 1995; Morishita & Arai, 1997; Obata et al., 1997).
The several types of mafic layers in the Horoman complex are usually concordant with the foliation of peridotite (Niida, 1984; Shiotani & Niida, 1997; Takazawa et al., 1999). It is difficult to estimate their P–T paths because of subsolidus re-equilibration at low pressures (Takazawa et al., 1999). Takazawa et al. (1999) suggested an origin by polybaric crystallization for the primary mineral assemblages of the several types of mafic layers based on detailed geochemical characteristics, including Sr, Nd and Pb isotopic ratios; one was initially garnet clinopyroxenite (Type I) and another was a plagioclase-rich cumulate (Type II). The nomenclature of the mafic layers in the Horoman complex will be discussed in detail in subsequent sections. Whole-rock Sm–Nd isotope data for the Type II mafic rocks are consistent with an 830 Ma whole-rock peridotite isochron (Yoshikawa & Nakamura, 2000) whereas the formation age of the Type I mafic rock is 80 Ma (Sm–Nd whole-rock isochron age) (Takazawa et al., 1999). From these lines of geochemical evidence, Takazawa et al. (1999) presented a hypothesis that the Type II mafic rocks experienced higher-pressure metamorphism after they had been formed from a melt as igneous cumulates. However, there have been no descriptions of petrographical characteristics indicating their high-pressure origin. One of the Type II rocks contains corundum as briefly reported by Morishita & Kodera (1998). They suggested that corundum could be formed by metamorphism of a plagioclase-rich protolith but did not present detailed descriptions of the corundum-bearing rock or its relationships with other corundum-free mafic layers. Corundum-bearing Type II mafic rocks may provide us with vitally important information on the tectonic history of the Horoman complex. In this paper we present detailed petrological data for the corundum-bearing mafic rock and related rocks, and discuss their petrogenesis and its significance for the tectonic history of the Horoman complex.
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GEOLOGICAL BACKGROUND |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Geological outline
The Horoman peridotite complex is located at the south end of the Hidaka metamorphic belt (e.g. Komatsu et al., 1986
). The Hidaka metamorphic belt consists of the Poroshiri ophiolite zone (Western Zone) of metamorphosed ophiolite (Miyashita, 1983; Arai & Miyashita, 1994) and the Main Zone of various kinds of plutonic and metamorphic rocks of various grades ranging from greenschist to granulite through amphibolite facies (Komatsu et al., 1986; Maeda et al., 1986; Osanai et al., 1986; Shiba, 1988; Toyoshima et al., 1994). The boundary between them is a large thrust fault (the Hidaka Main Thrust) associated with mylonite over the whole extent of the metamorphic belt (e.g. Komatsu et al., 1986). Several tectonic sheets of upper-mantle peridotite are inserted in the basal mylonite zone of the Main Zone (Niida & Katoh, 1978; Komatsu et al., 1994). The Horoman peridotite complex is the largest one, 8 km x 10 km in area and >3 km in thickness (e.g. Igi, 1953; Komatsu & Nochi, 1966; Niida, 1984; Sawaguchi & Takagi, 1997; Sawaguchi, 1999). We investigated the southern part of the complex (Fig. 1). The complex consists of layered peridotite with small amounts of mafic rocks and is divided into two parts: the Upper and Lower Zones (Komatsu & Nochi, 1966; Niida, 1984) (Fig. 1). The Upper Zone is characterized by an abundance of mafic layers and by sharp lithological boundaries on a scale of a few centimeters. The Lower Zone is characterized by gradational lithological boundaries (Komatsu & Nochi, 1966; Niida, 1984).
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Ozawa & Takahashi (1995) examined systematic changes in mineralogical and petrographical features from the Upper to Lower Zones and chemical zoning of minerals from the plagioclase peridotite. They suggested that the Horoman peridotite ascended as a mantle diapir from the garnet peridotite stability field, following a higher-temperature decompression path (1150°C at >2·0 GPa) for the Upper Zone and a lower-temperature path (900°C at >2·0 GPa) for the Lower Zone. Ozawa (1997) suggested another possibility, that the Horoman peridotite experienced a heating event at a pressure near the boundary between the spinel and plagioclase peridotite facies during the decompression process.
Occurrence and classification of peridotites
On the basis of petrography and mineralogy, the peridotites are divided into three suites (Takahashi, 1991, 1992): (1) the Main Harzburgite–Lherzolite suite (MHL); (2) the Spinel-rich Dunite–Wehrlite suite (SDW); (3) the Banded Dunite–Harzburgite suite (BDH). The main part of the complex is composed of the MHL, i.e. harzburgite, spinel lherzolite and plagioclase lherzolite with gradational lithological boundaries. This suite has residual characteristics indicating various degrees of melt extraction (Obata & Nagahara, 1987; Frey et al., 1991; Takahashi, 1991, 1992; Takazawa et al., 1992, 2000; Yoshikawa & Nakamura, 2000). The SDW suite occurs in the MHL harzburgite as thin layers (up to 15 m) (Takahashi, 1992, 1997; Fig. 1). The SDW suite consists of dunite and wehrlite containing abundant large chromian spinel grains, and is a cumulate from a melt genetically related to the MHL suite (Takahashi, 1991, 1997; Yoshida & Takahashi, 1997). The BDH suite consists of conspicuously layered dunite, harzburgite and olivine orthopyroxenite with sharp lithological boundaries, and is a cumulate from a high-Mg, high Cr/Al melt such as high-Mg andesite (Takahashi, 1991, 1992). The BDH suite has been interpreted as exotic blocks captured by the MHL suite (Takahashi, 1991, 1992).
Classification of mafic rocks
The mafic rocks in the Horoman complex have been divided into several types (Niida, 1984; Shiotani & Niida, 1997; Takazawa et al., 1999). Niida (1984) first suggested the existence of two types of mafic rock, GB I and II. Shiotani & Niida (1997) extended this work and divided the mafic rocks into four types (GB I to GB IV) on the basis of mineral assemblages and chemical compositions. GB I, which is the most abundant type, is characterized by the presence of Ti-rich minerals such as titaniferous clinopyroxene, kaersutite and ilmenite. GB II, which is the second most abundant type, is marked by the presence of low-Ti clinopyroxene. Both GB I and II are poor in orthopyroxene, and contain brown and green spinels as accessory minerals. GB III contains clinopyroxene, orthopyroxene and spinel. GB IV contains abundant orthopyroxene and spinel. Takazawa et al. (1999) basically followed Shiotani & Niida’s discrimination for mafic layers but they renamed the GB I to GB IV types of Shiotani & Niida (1997) as Type I to IV mafic granulites, respectively. Takazawa et al. (1999) also added a new type of mafic layer, Type V mafic granulite, which occurs in harzburgite near the BDH suite and has a recrystallized mineral assemblage consisting of orthopyroxene, green spinel, plagioclase, clinopyroxene and minor amounts of olivine and phlogopite.
In this paper, we basically follow the nomenclature of Takazawa et al. (1999) and call the mafic rocks Types I to V. Corundum-bearing mafic rocks belong to the Type II mafic rock based on petrography and mineral chemistry (Morishita & Kodera, 1998). We, therefore, focus on the Type II mafic rocks (± corundum) and related rocks.
Layered structure of the Horoman complex: a symmetrical feature
The feature particular to the Horoman complex is a symmetrical layered sequence consisting of the MHL peridotites (plagioclase peridotite, symplectite-bearing spinel peridotite and spinel peridotite of Fig. 1) and the minor distinctive SDW peridotites (Komatsu & Nochi, 1966; Niida, 1984) (Fig. 2). This symmetrical layered sequence is also distinctive in terms of both bulk-rock and mineral chemistries (Niida, 1984; Obata & Nagahara, 1987; Takahashi, 1991, 1992; Takazawa et al., 1992), and it repeats several times with intervals of a few meters to a few hundred meters in the complex (Fig. 1). Takahashi (1991, 1992) emphasized that the symmetrical layered unit is basically characterized by an arrangement of a cumulus peridotite (SDW) in the middle of a series of residual peridotites (MHL) in which refractory, residual character increases (plagioclase peridotite to spinel peridotite layers, Fig. 1) toward the cumulus peridotite (SWD) (Fig. 2). The Type II mafic rocks occur in the SDW suite as minor components of the symmetrical layered structure (Figs 1 and 2).
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PETROGRAPHY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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The petrography of the peridotites and mafic rocks in the Horoman complex has been described in detail by previous workers (e.g. Niida, 1984
; Takahashi, 1991; Shiotani & Niida, 1997; Takazawa et al., 1999). The Type II mafic rocks are summarized below and the corundum-bearing rock and related rocks are described in more detail.
Ordinary Type II mafic rocks
The Type II mafic rocks are easily distinguished from other mafic rocks by their light green to blue color in outcrop (Takazawa et al., 1999). It is noteworthy that the Type II mafic rocks appear exclusively in the SDW suite (Shiotani & Niida, 1997; Takazawa et al., 1999) (Figs 1 and 2). They occur as thin layers (1–200 cm in thickness), which are concordant both with the foliation of the peridotite (SDW) and with the lithological boundary planes between the SDW and MHL suites (Fig. 3a and b). Some layers show isoclinal folding, boudinage and slump-like structures (M. Obata, personal communication, 1996; Takazawa et al., 1999). Dunite layers with plagioclase-rich seams characteristically alternate with the SDW peridotites and Type II mafic layers on a scale of a few tens of centimeters in the Upper Zone (Figs 2b and 3b). The Type II mafic rocks are more common in the Upper Zone than in the Lower Zone (Fig. 1). The SDW peridotite layer occurring with the Type II mafic rocks is traceable over at least 4 km along strike in the Upper Zone (Fig. 1). In contrast, we have confirmed that the thickest Type II mafic rock is traceable only for a few tens of meters in the Lower Zone (B250 of Fig. 1). The contact between the Type II mafic rocks and SDW peridotite is usually very sharp (Fig. 3a). Modal layering is usually visible on both weathered and sawn surfaces of the Type II mafic rock. We have found that several coarse spinel grains (up to 3 cm across) make concordant trails within one thick layer at only one outcrop (62501 of Fig. 1) (Fig. 3c).
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The Type II mafic rocks show fine-grained (<0·5 mm across) equigranular metamorphic textures consisting of plagioclase, clinopyroxene and olivine with small amounts of amphibole, green and brown spinels and orthopyroxene, but quartz is absent (Fig. 3d) (Niida, 1984
; Shiotani & Niida, 1997; Takazawa et al., 1999). Brown and green spinels associated with discrete plagioclase are relatively small in size, <0·3 mm across. Relatively large green spinel grains, up to 3 cm across, have ragged rims surrounded by plagioclase rinds. In some cases, symplectic texture, possibly consisting of Al-spinel and amphibole, occurs at grain boundaries between olivine and plagioclase (Morishita, 1999). Because it is different from the symplectite of possible garnet origin (e.g. Kushiro & Yoder, 1966; Takahashi & Arai, 1989; Ozawa & Takahashi, 1995; Morishita & Arai, 1997; Obata et al., 1997), we call it ‘micro-symplectite’ hereafter.
In the case of the Lower Zone, the Type II mafic layer, 2 m in thickness, has seams enriched with symplectite and fine-grained aggregates consisting of plagioclase, spinel and orthopyroxene at its outermost margins at the locality B250 (Fig. 1) (Morishita & Arai, 1997; Fig. 3e and f). Lenticular aggregates of clinopyroxene, of 2 mm grain size, with orthopyroxene lamellae occur in some symplectite-rich layers. SDW peridotite occurs as thin (4 cm in thickness) layers along the margins of the Type II mafic rock.
A dunite layer with plagioclase-rich seams shows an equigranular texture consisting of relatively fine-grained olivine neoblasts (<2 mm across) with rare porphyroclasts (up to 5 mm across). The dunite layer is similar in texture to the SDW peridotite, but the discrete brown spinel is less abundant in the former than in the latter. The plagioclase-rich seam consists of fine-grained plagioclase, olivine and brown spinel (<0·5 mm across) (Fig. 3g). Aggregates of small spinel grains (<0·1 mm across) and plagioclase in the plagioclase-rich seam (Fig. 3g) are similar to the symplectite in the MHL peridotite. Discrete brown spinel, up to 3 mm across, has a ragged rim surrounded by a plagioclase rind.
Corundum-bearing Type II mafic rock
Type II mafic rocks with corundum show a roughly layered structure consisting of olivine-rich and olivine-poor parts (Fig. 4a–c). Corundum occurs in relatively thick olivine-poor lenticular portions, which are almost concordant with the fine layered structure (Fig. 4a and b). Coarse-grained spinel (up to 5 cm across) and lenticular spinel–plagioclase aggregates occur sporadically in the corundum-bearing portion (Fig. 4a and b). We can recognize strong deformation of the corundum-bearing part by elongation of the lenticular spinel–plagioclase aggregate, which is concordant with the foliation plane of the other part (Fig. 4b).
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A zonal arrangement of inner green spinel and outer plagioclase is observed around pink corundum (up to 3 cm across) (Fig. 4a, b, d and e). We refer to this zone around corundum as the plagioclase–spinel zone hereafter. Plagioclase around the plagioclase–spinel zone is mostly saussuritized and in some cases is in contact with corundum. Margarite occurs at boundaries between corundum and plagioclase of the plagioclase–spinel zone. Very fine-grained aggregates of brown spinel, <35 µm across, very rarely occur in the plagioclase–spinel zone and are sometimes associated with clinopyroxene (Fig. 4f). Rarely corundum includes saussuritized plagioclase (Fig. 4d). One inclusion consists of several grains of plagioclase showing equigranular texture (Fig. 4g) in which the cores were susceptible to partial alteration. Green spinel and margarite occur in some cases at the boundary between corundum and the plagioclase inclusion (Fig. 4d and g). Very fine-grained brown spinels, <20 µm across, very rarely occur as an inclusion in corundum. Quartz has not been detected as an inclusion. Fine-grained green spinel, <0·2 mm across, is commonly present in the olivine-poor layer but may be concentrated to form green spinel-rich lenticular aggregates (Fig. 4h). In the olivine-rich layer, brown spinel, which is extremely rare (<0·1% in mode) and fine grained (<0·1 mm across), is discrete and associated with plagioclase in the olivine-rich layer. Amphibole is more abundant in the olivine-rich layer than in the ordinary corundum-free Type II mafic layers. The micro-symplectite is also well developed in the olivine-rich layer.
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CHEMISTRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Whole-rock chemical compositions
Whole-rock major element chemical compositions of the Type II mafic rocks (± corundum) and related dunite layers with plagioclase-rich seams were determined on fused disks using a Rigaku (System 3270) X-ray fluorescence spectrometer at Kanazawa University (Table 1). The analysis was performed using an accelerating voltage of 50 kV and a beam current of 20 mA. It is difficult to directly determine the bulk chemical compositions for the corundum-bearing part of a sample because these tend to be very heterogeneous as a result of the layered structure and sporadic occurrence of large mineral clots (corundum + green spinel + plagioclase). We prepared corundum-bearing rocks including and without large corundum–spinel–plagioclase clots (both >100 g) as follows. The sample was sliced into thin slabs with a diamond saw to obtain the corundum-bearing rock including large mineral clots. The composition of the corundum-free parts was obtained by eliminating corundum–spinel–plagioclase clots from the slabs using a diamond saw. We calculated the CIPW norm for these rocks and confirmed that they are extremely low in Si, and Si-undersaturated minerals such as kalsilite, Ca-orthosilicate and wüstite appear in the corundum-bearing rocks (Table 1).
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CaO and Al2O3 contents (wt %) vary linearly against MgO and SiO2 contents (wt %) (Fig. 5). The compositions of the corundum-bearing rock plot at the MgO-poor, SiO2-rich ends of the chemical trends of ordinary Type II mafic rocks whereas dunite layers with plagioclase-rich seams plot at the MgO-rich, SiO2-poor ends.
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The Type II mafic rocks are characterized by low TiO2 (Shiotani & Niida, 1997), high Al2O3 and high mg-number [100Mg/(Mg + Fetotal) atomic ratio] compared with the Type I mafic rocks (Takazawa et al., 1999). Similarly, two types of mafic rocks are recognized in garnet-bearing pyroxenites (and granulite) from both the Beni Bousera and Ronda massifs (Kornprobst et al., 1990; Morishita, 1999; Takazawa et al., 1999). Although the TiO2 contents of almost all the samples are below the detection limit (<0·07 wt %), the Type II mafic rocks are slightly lower in mg-number and higher in TiO2 (0·1 wt %) in the Lower Zone than in the Upper Zone.
Mineral chemical compositions
Mineral chemical compositions were determined with a JEOL 8800R electron microprobe at Kanazawa University, using the ZAF correction method. Typical operating conditions were 15 keV accelerating potential, 12 nA beam current measured on Faraday cup and 3 µm beam diameter for major elements; and 25 keV accelerating potential, 20 nA beam current measured on Faraday cup and 3 µm beam diameter for NiO measurement on olivine. Natural and synthetic minerals were used for standards. Representative analyses of the rock-forming minerals are shown in Table 2.
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Ordinary Type II mafic rocks free of corundum
Forsterite and NiO contents of olivine in the ordinary Type II mafic rocks of the Upper Zone are 86–89 and 0·21–0·25 wt %, respectively (Fig. 6). Clinopyroxenes are poor in TiO2 (<0·4 wt %) and Cr2O3 (<0·8 wt %) (Fig. 7). Anorthite contents [100 Ca/(Ca + Na) atomic ratio] of plagioclase range from 83 to 91. The cr-numbers [Cr/(Cr + Al) atomic ratios] in green and brown spinels are <0·1 and >0·25, respectively.
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Type II mafic rocks in the Lower Zone (B250 section) show grain-by-grain chemical heterogeneity (Figs 6 and 7). A systematic variation of mineral chemistry is observed, however, across the lithological boundaries as follows (Morishita & Arai, 1997). Going from the symplectite-bearing margin of the mafic layers into the middle of the layer, the forsterite content of the olivine decreases from 91 to 82, the anorthite content of plagioclase decreases from 91 to 79 and the cr-number of the spinel decreases from 0·30 to 0·05.
Corundum-bearing Type II mafic rocks
The mg-number [Mg/(Mg + Fe) atomic ratio] of clinopyroxene is high (0·93) in the corundum-bearing rocks but relatively low (0·86) in the olivine-rich rocks that are intercalated with the corundum-bearing rocks (Fig. 7). The high mg-number of clinopyroxene in the corundum-bearing rocks may be due to subsolidus reaction with spinel (mg-number is 0.82). Clinopyroxenes are poor in TiO2 (<0·2 wt %) and Cr2O3 (<0·6 wt %) (Fig. 7). Clinopyroxene associated with very fine-grained brown spinels in the plagioclase–spinel zone around corundum is higher in Cr2O3 (>1 wt %, cr-number 0·54) than clinopyroxene in the matrix.
The anorthite content of plagioclase in the Type II mafic rocks intercalated with corundum-bearing rocks is usually 89, but more calcic (>96) in a partly saussuritized area. Plagioclase is free of chemical zoning in the matrix of the corundum-bearing rocks whereas some plagioclase occurring as inclusions within corundum (Fig. 4g) shows chemical zoning, i.e. from An72 at the core to An96 at the rim. Plagioclase with an anorthite content as low as An55 has been detected as inclusions in corundum. The anorthite content of plagioclase increases toward green spinel in the plagioclase–spinel zone around corundum.
The cr-numbers in green and brown spinels are <0·1 and >0·5, respectively. Very fine-grained spinel occurring as inclusions in corundum and in the plagioclase–spinel zone has high cr-numbers, i.e. 0·6–0·7. The forsterite and NiO contents of olivine in the olivine-rich rocks intercalated with corundum-bearing rocks are 86 and 0·14 wt %, respectively (Fig. 6). Amphibole is pargasite (Leake, 1978) and is also poor in TiO2 (<0·6 wt %). Margarite is low in paragonite (<10 mol %) and muscovite (almostly nil) components. Corundum is almost pure Al2O3 and has low FeO (<0·3 wt %), Cr2O3 (<0·2 wt %) and TiO2 (<0·1 wt %) contents. We need a more careful study of the chemistry of corundum because trace amounts of Fe, Cr, Ti and Ga in corundum may be very significant in interpreting the origin of corundum (Sutherland et al., 1998).
Dunite layers with plagioclase-rich seams
Olivine in the dunite layers with plagioclase-rich seams plots in the low Fo–NiO part of the SDW olivine field (Fig. 6). TiO2 and Cr2O3 contents of clinopyroxene are <0·3 wt % and <1 wt %, respectively (Fig. 7). The anorthite content of plagioclase is 90. The cr-number is 0·2 in coarse-grained spinel core but >0·3 in fine-grained spinel associated with plagioclase.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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Origin of the Type II mafic rocks
Whole-rock and mineral chemical compositions of the Type II mafic rocks (± corundum) and dunite layers with plagioclase-rich seams form linear chemical trends in oxide co-variation diagrams (Fig. 5). The corundum-bearing rocks and dunite layers with plagioclase-rich seams plot on the Al-rich and Al-poor extensions, respectively, of chemical trends defined by the corundum-free Type II mafic rocks. The corundum-bearing Type II mafic rocks are characterized by high Al2O3, CaO and SiO2, and low MgO contents. Mechanical mixing of corundum by deformation or as corundum megacrysts in basalt dykes (Sutherland et al., 1998
) cannot explain the Ca enrichment in the corundum-bearing rocks. Both the high Ca and high Al are interpreted as intrinsic properties of the protolith of the Type II mafic rocks, and an exotic or xenocrystal origin for corundum is improbable.
Bulk chemical compositions of the Type II mafic rocks and dunite layers with plagioclase-rich seams were projected from clinopyroxene onto the plane olivine (M2S)–Ca-Tschermak pyroxene (CaTs; CAS)–quartz (S) and from CaTs onto the plane clinopyroxene (CMS)–olivine (M2S)–quartz (S) by the method of O’Hara (1968) because of some difficulties with the CIPW norm for these rocks as suggested previously. These define a line on the olivine–clinopyroxene–anorthite plane (Fig. 8), suggesting that they could be mixtures of plagioclase and olivine with minor amounts of clinopyroxene and spinel, i.e. their troctolitic to olivine gabbroic protoliths formed at low pressure by simultaneous precipitation of plagioclase and olivine. This condition is <1 GPa in the system CaO–MgO–Al2O3–SiO2 (CMAS) (e.g. Presnall et al., 1978), corresponding to the base of the crust or uppermost mantle. We therefore conclude that the Type II mafic rocks (± corundum) and dunite layers with plagioclase-rich seams were originally a series of cumulus rocks formed by crystallization of olivine, plagioclase and clinopyroxene; the corundum-bearing rocks were derived from plagioclase-rich varieties, and the dunite layers with plagioclase-rich seams from olivine-rich varieties.
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The hypothesis that protoliths of the Type II mafic rocks formed at low-pressure conditions is consistent with Shiotani & Niida (1997) and Takazawa et al. (1999). Takazawa et al. (1999) showed that the Type II mafic rocks have positive anomalies of Sr and Eu, suggesting plagioclase crystallization in their genesis. Furthermore, Shiotani & Niida (1997) and Takazawa et al. (1999) demonstrated that the Type II mafic rocks are similar in their bulk 87Sr/86Sr isotope ratio to N-MORB and considered that the Type II mafic rocks were formed from MORB or a MORB-like magma.
Relationships between peridotites (the SDW and MHL suites) and Type II mafic rocks
As suggested above, one of the features of the Horoman complex is the repeated symmetrical layered structure. Takahashi (1991, 1992) explained the layered structure by melting, melt extraction and crystallization along cracks in the mantle. On the other hand, Toramaru (1997) and Toramaru et al. (2001) pointed out that the symmetrical layer structure was formed by stretching and folding of a simple pair of residue–cumulate in the mantle. Takazawa et al. (2000) and Yoshikawa & Nakamura (2000) also ascribed the compositional layering in the peridotite to thinning and folding at subsolidus conditions of a large-scale simply layered body that had been formed by polybaric melt extraction. The physical juxtaposition of the SDW suite and the MHL suite may be related to the partial melting event, which caused a decrease of melt components in the MHL suite towards the SDW suite (Takahashi, 1991, 1992).
Both Type II mafic rocks and dunite layers with plagioclase-rich seams exclusively occur in the SDW peridotite in the middle of the symmetrical layered structure (Fig. 2). It is difficult, therefore, to consider that the Type II mafic rock was selectively formed only in the SDW peridotite after formation of the symmetrical layered structure of peridotite. Whole-rock and mineral compositions gradually change from the SDW peridotites to Type II mafic rocks through to dunite layers with plagioclase-rich seams (Shiotani & Niida, 1997). These observations suggest that the Type II mafic rocks and dunite layers with plagioclase-rich seams were crystallized from the melt that was responsible for the formation of the SDW cumulus peridotite, i.e. during the major melting event that occurred in the MHL suite.
Petrogenesis of corundum-bearing Type II mafic rocks
Corundum disappearance
Textural relationships of corundum-bearing Type II mafic rocks indicate that the corundum incompletely reacted with other phases into green spinel and plagioclase, thereby showing that corundum was not stable in the Type II mafic rocks during later P–T conditions of the Horoman complex. Although corundum has not been found in the Type II mafic rocks from outcrops, the presence in one outcrop of coarse-grained green spinel, up to 2 cm, surrounded by plagioclase (Figs 1 and 3c) may indicate completion of the corundum disappearance reaction. In the CMAS system, a possible reaction for corundum disappearance is given by end-members of minerals (e.g. Nozaka, 1997; Morishita & Kodera, 1998):
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) and the thermodynamic data of Holland & Powell (1990) (Fig. 9). It should be noted that the presence of Fe in the system will stabilize the spinel, which accommodates iron preferentially and therefore shifts the reaction to lower pressure and temperature. On the other hand, the presence of Na in the system will stabilize the plagioclase at higher pressure. Furthermore, reaction (1) will not occur at >2 GPa because of the disappearance of plagioclase. Despite these problems, this reaction may proceed to the right by heating and/or decompression (Morishita & Kodera, 1998) (Fig. 9).
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To determine the P–T path for the corundum disappearance reaction in the Horoman complex, we estimated the latest P–T history of the Type II mafic rocks in the plagioclase–olivine stability field. We obtained >1000°C (5a in Fig. 9) as the equilibrium temperature using Mg–Fe2+ partitioning between olivine and clinopyroxene rims or fine-grained clinopyroxenes in the Type II mafic rocks and dunite layers with plagioclase-rich seams (Hiramatsu & Hirajima, 1995). Under such conditions, margarite may be formed by the reaction (e.g. Chatterjee, 1976; Perkins et al., 1980
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) (5b in Fig. 9). Ozawa & Takahashi (1995) and Takazawa et al. (1996) estimated the temperature as 750–850°C at 0·5 GPa for the rims of pyroxenes in plagioclase lherzolite. These observations indicate that Type II mafic rocks, as well as peridotite, recorded cooling within the plagioclase–olivine stability field (broken arrow from 5a to 5b in Fig. 9). It has been considered that the Horoman plagioclase peridotite experienced nearly adiabatic decompression during upwelling from the garnet to plagioclase peridotite stability fields followed by rapid cooling in the plagioclase peridotite stability field (Ozawa & Takahashi, 1995; Takazawa et al., 1996; Ozawa, 1997; Yoshikawa & Nakamura, 2000). Thus the corundum disappearance was possibly caused by decompression as a result of uplift of the Horoman complex (4 in Fig. 9).
Corundum formation
Two possibilities can be considered for the genesis of corundum in Type II mafic rocks: (1) as a residue after partial melting; (2) metamorphism of the Type II protolith.
Residue from partial melting. Råheim & Green (1974) experimentally determined the solidus and liquidus relations for the Lunar Highland Basalt composition (gabbroic anorthosite), which is very similar to that of the corundum-bearing rocks from the Horoman complex (Fig. 5). They confirmed that corundum appeared as one of the liquidus phases crystallizing from melt with clinopyroxene and spinel or with clinopyroxene and garnet at P > 1·8 GPa, T > 1400°C (3' in Fig. 9).
This model requires two steps for corundum formation in the Type II mafic rocks: (1) formation of high-Al mafic compositions, that is, the Type II protolith formed at pressures <1·0 GPa (an assemblage of olivine + plagioclase + minor clinopyroxene + minor spinel); (2) later partial melting of the Type II protolith at higher pressure, >1·8 GPa.
If the corundum was formed in this way, subsolidus recrystallization of the corundum-bearing residues would erase their igneous textures, in which case we need to search for evidence for the igneous origin of corundum in the geochemical characteristics of the Type II mafic rocks. This will be discussed in a separate paper.
Metamorphic origin. Experimental results of Råheim & Green (1974) showed that there was no corundum at subsolidus conditions in their high-Al composition. However, their results cannot be directly applied to the studied rocks because the mg-number and Ca and Na contents, which are critical to the compositions of garnet and clinopyroxene, are higher in the corundum-bearing mafic rocks from the Horoman complex than those found by Råheim & Green (1974). Thus we infer that corundum can form as one of the subsolidus phases in the studied high-Al2O3 composition and discuss a metamorphic origin for corundum. Corundum can be formed by reaction (1), that is spinel + anorthite = 2 corundum + diopside, in the Type II mafic rocks with sufficiently high Al2O3 contents, e.g. a troctolitic protolith. This reaction needs spinel as one of the reactants, which may be formed by the reaction (e.g. Kushiro & Yoder, 1966)
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). As discussed above, corundum disappearance was possibly caused by decompression at high temperature (>1000°C) as a result of uplift of the Horoman complex. It is, therefore, difficult to consider that corundum disappearance occurred after corundum formation by isobaric cooling. We conclude that corundum was formed by metamorphism at higher pressure (>1·5 GPa at 1000°C) after the formation of Type II mafic rocks as igneous cumulates at shallow level (<1·0 GPa) (3 in Fig. 9).
If this is true, the next question is the maximum P–T condition recorded in the Type II mafic rock. Kyanite has not been observed in Type II mafic rocks and their low SiO2 content is consistent with this. We have observed symplectites of pyroxene, spinel and plagioclase in the Type II mafic rocks and dunite layers with plagioclase-rich seams. These symplectites are similar to those in the MHL peridotite and have been interpreted to be of pyrope-rich garnet origin (e.g. Kushiro & Yoder, 1966; Takahashi & Arai, 1989; Ozawa & Takahashi, 1995; Obata et al., 1997). As will be discussed in a separate paper, the presence of symplectites suggests that the Type II mafic rocks were possibly metamorphosed from garnet-bearing pyroxenite with decreasing pressure.
Corundum-bearing garnet pyroxenites alternating with peridotites have been reported from some peridotite complexes of high-pressure origin, e.g. the Beni Bousera in Northern Morocco (Kornprobst et al., 1990), the Ronda in Southern Spain (Morishita, 1999; Sánchez-Rodríguez & Gebauer, 2000; Morishita et al., 2001), Cabo Ortegal massif in northwestern Spain (Girardeau & Gil Ibarguchi, 1991) and the Val Malenco in the Italian Alps (Müntener & Hermann, 1996). In the Ronda, although corundum-bearing rocks are found in garnet-bearing peridotites, the garnets in corundum-bearing rocks have selectively suffered kelyphitization because they are richer in pyrope component than corundum-free rocks, corresponding to their high-mg-number bulk composition (Morishita et al., 2001). Thus, garnets in mafic rocks of the Horoman complex could completely disappear during the latest P–T history because of the high temperatures maintained during upwelling to the plagioclase peridotite stability field. Corundum-bearing eclogites (grospydites) have also been reported as xenoliths in kimberlite pipes (e.g. Sobolev et al., 1968; Ater et al., 1984). Corundum coexisting with both garnet and clinopyroxene also occurs as inclusions in diamond (Sobolev et al., 1998). It is difficult to estimate precise P–T conditions at which corundum is stable in equilibrium with garnet and clinopyroxene because corundum would be stable in garnet pyroxenites at a wide range of pressure (e.g. Gasparik, 1984a).
Implication for the recycling of the Horoman peridotite
Both origins for corundum, i.e. as a residue after partial melting and metamorphism, require two steps for corundum formation in the Type II mafic rocks: (1) formation of high-Al mafic compositions, i.e. the Type II protolith formed at pressures <1 GPa; (2) later high-pressure metamorphism (P > 1·5 GPa) and/or partial melting of the Type II protolith at higher pressure, >1·8 GPa. Both possibilities for the origin of corundum are consistent with the complex P–T history for the Type II mafic rocks proposed by Takazawa et al. (1999) based on geochemical and chronological data, i.e. high-pressure recrystallization after initial formation at lower-pressure conditions. However, the residue from partial melting requires a higher temperature in the early stages of upwelling of the Horoman peridotite than suggested by Ozawa & Takahashi (1995).
Petrological data in this paper, combined with other geochemical and chronological data from the literature, possibly indicate a complicated spiral-like P–T trajectory for the Horoman peridotite. This could be interpreted as evidence for multiple recycling of mantle peridotite that has ascended and descended more than once.
Yoshikawa & Nakamura (2000) showed that the Horoman peridotite experienced polybaric melt extraction within the garnet to spinel peridotite stability field at 833 ± 78 Ma, estimated by Sm–Nd isotope systematics of whole-rock samples. Takazawa et al. (2000) concluded on the basis of geochemical data that the Horoman peridotite is the residue after partial melting of garnet peridotite. Takazawa et al. (1999) suggested, from the Sm–Nd isochron diagram for the Type II mafic rocks combined with peridotites, that the Type II mafic rocks were formed shortly after initial partial melting of the Horoman peridotite at 830 Ma. This age is much older than the 80 Ma age obtained from an Sm–Nd isochron for Type I mafic layers.
Takazawa et al. (2000) considered that the Horoman peridotite was uplifted and emplaced into the crust under subsolidus conditions after a partial melting event, following the P–T paths suggested by Ozawa & Takahashi (1995) and Takazawa et al. (1996). However, if the Type II mafic rock experienced high-pressure metamorphism after it had formed as an igneous cumulate at low pressures as a member of the symmetrical layers, the peridotite should have also experienced the same high-pressure metamorphism.
A complex P–T trajectory is, therefore, proposed for the whole Horoman complex (Fig. 9). Previous workers (Takazawa et al., 2000; Yoshikawa & Nakamura, 2000) have argued that the Horoman complex started to melt within the garnet peridotite stability field and lost melt fractions in the garnet to spinel peridotite fields at 830 Ma (1 in Fig. 9). The residue ascended to the plagioclase–olivine stability field, retaining a small amount of melt within the upwelling body. Separation of this melt and partial crystallization of liquidus phases from the melt formed a cumulus rock series, i.e. the SDW suite and Type II protolith, at low pressures within a short interval (a few million years?) after partial melting and upwelling of the peridotite (2 in Fig. 9). After cooling, the Horoman peridotites and Type II protoliths experienced higher-pressure metamorphism (>1·5 GPa) (3 in Fig. 9), which was probably caused by subduction. It is possible that the Type II mafic rocks experienced partial melting within the garnet stability field (3' in Fig. 9). Finally, the Horoman complex ascended from the garnet to plagioclase peridotite stability field again (4 in Fig. 9), following the decompression P–T path suggested by Ozawa & Takahashi (1995) and Takazawa et al. (1996). The protolith of Type I mafic rock, i.e. garnet pyroxenite, intruded into peridotites at 80 Ma (Takazawa et al., 1999). The results of this study cannot constrain whether this occurred during compression or decompression. To prove this suggestion, we should determine the highest P–T regime experienced by the Type II mafic rock. Until then, the P–T–t relationship between peridotite and mafic layers will remain an important question in the Horoman complex.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDPETROGRAPHYCHEMISTRYDISCUSSIONCONCLUSIONSREFERENCES |
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- Only corundum-bearing mafic rocks [referred to by Takazawa et al. (1999) as Type II mafic rocks] appear to preserve the relics of a high-pressure mineralogy (>1·5 GPa), although almost all mafic layers in the Horoman complex at present record an approach to equilibrium within the plagioclase peridotite stability field (Takazawa et al., 1999). These rocks occur as alternating layers with the cumulus peridotite [the SDW suite of Takahashi (1991, 1992)] and are minor components of the symmetrical layered structure.
- Whole-rock and mineral chemical compositions of the Type II mafic rocks show their cumulus origin, with a primary mineral assemblage controlled by crystallization of olivine, plagioclase and clinopyroxene indicative of a low-pressure origin at <1·0 GPa. The corundum-bearing rocks were derived from plagioclase-rich varieties, whereas the dunite layers with plagioclase-rich seams were from olivine-rich varieties.
- On the basis of mineralogy the Type II mafic rocks cooled from 1000°C to 600°C within the plagioclase–olivine stability field on the latest P–T path, which is similar to that recorded by the peridotite (Ozawa & Takahashi, 1995; Takazawa et al., 1996). The corona texture around the relic corundum suggests that corundum was not in equilibrium with other minerals during the latest P–T condition of the Horoman complex. Corundum disappearance was probably caused by decompression by reaction with clinopyroxene during ascent of the Horoman complex.
- There are two possibilities for the formation of corundum in the Type II mafic rocks. One is that corundum formed by metamorphism of the Type II protolith at higher-pressure conditions (>1·5 GPa at 1000°C) than that of its formation as an igneous cumulate. The whole-rock composition would then control the presence or absence of corundum. The other is that corundum formed as a residual phase during partial melting of the Type II protolith at >1·8 GPa.
- Field and petrographical observations coupled with the isotopic date (Takazawa et al., 1999; Yoshikawa & Nakamura, 2000) suggest that the Type II mafic protolith (i.e. low-pressure olivine–plagioclase cumulates) formed within the peridotite at the time of initial formation of the Horoman peridotite as a residue, emplaced at shallow level, from primitive mantle at 830 Ma. After cooling at shallow levels, the Type II mafic rocks, as well as surrounding peridotites, experienced subduction to the garnet stability field. Some of the Type II mafic rocks may have experienced partial melting within the garnet stability field. Finally, the Horoman complex ascended from the garnet to plagioclase peridotite stability field following the P–T path of peridotite suggested by Ozawa & Takahashi (1995) and Takazawa et al. (1996).
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ACKNOWLEDGEMENTS |
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We thank Harry Becker, Jon Snow, Eiichi Takazawa and anonymous reviewer for their constructive comments. We also thank Pamela D. Kempton for her editorial advice. We are grateful to David H. Green for discussions about the origin of corundum in mafic rocks, and to Hiroshi Shukuno, Tatsuki Tsujimori, Akira Ishiwatari, Yoshihiko Tamura and Atsushi Toramaru for their daily discussion and technical help with EPMA and XRF analyses. Tadahiro Kodera helped us to collect the samples. We thank the Board of Education of Samani Town for kindly permitting use of the facilities of the Managing Office of Apoi-dake to make some thin sections. This research was partly supported by the JSPS Fellowships for Japanese Junior Scientists (T.M.).
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FOOTNOTES |
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*Corresponding author. Telephone: +81-76-264-5723. Fax: +81-76-264-5746. E-mail: moripta{at}kenroku.kanazawa-u.ac.jp
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