Insight into the uppermost mantle section of a maturing arc: The Eastern Mirdita ophiolite, Albania

Lithos 124 (2011) 215–226

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Lithos j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / l i t h o s

Insight into the uppermost mantle section of a maturing arc: The Eastern Mirdita ophiolite, Albania Tomoaki Morishita a,b,⁎, Yildirim Dilek c, Minella Shallo d, Akihiro Tamura a, Shoji Arai e a

Frontier Science Organization, Kanazawa University, Kanazawa 920-1192, Japan Department of Geology & Geophysics, University of Hawaii at Manoa, 1680 East-West Rd., Honolulu, HI 96822, USA Department of Geology, Miami University, Oxford, OH 45056, USA d Fakulteti i Gjeologise dhe Minierave, Universiteti Politeknik, Tirana, Albania e Department of Earth Sciences, Kanazawa University, Kanazawa 920-1192, Japan b c

a r t i c l e

i n f o

Article history: Received 14 March 2010 Accepted 4 October 2010 Available online 20 October 2010 Keywords: Ophiolite Mantle section Albania Mirdita Harzburgite–dunite–orthopyroxenite Island arc Mid-ocean ridge

a b s t r a c t We examined peridotite massifs in the eastern part of the Mirdita ophiolite (EMO), Albania, where arc-related magmas are abundant in the upper volcanic sequences. Structurally, clinopyroxene porphyroclast-bearing harzburgites (Cpx-harzburgite hereafter) occur in the lower parts of the peridotite massifs, whereas harzburgites and dunites are more abundant towards the upper parts. Dunite is commonly associated with chromitite layers. Orthopyroxenite occurs as dikes and/or networks at all structural levels, although it is more abundant in the uppermost sections. Orthopyroxenite commonly crosscuts the foliation of peridotites and the lithological boundaries between dunites (chromitite) and harzburgites, suggesting that it was formed in the late stage. Major and trace element compositions of minerals in the Cpx-harzburgites indicate that they were formed as the residue of less-flux partial melting, and are similar to those in abyssal peridotites from mid-ocean ridge systems. Harzburgites have more depleted major element compositions than the Cpx-harzburgites. Light rare earth element (LREE)-enrichment in clinopyroxene coupled with hydrous silicate mineral inclusions in spinels in harzburgites indicate that harzburgites were produced as a result of enhanced partial melting of depleted peridotites due to infiltration of hydrous LREE-enriched fluids/melts. Based on olivine and spinel chemistries, dunites are classified into two types: high-Cr# (= Cr/(Cr + Al) atomic ratio) spinel-bearing dunite and medium-Cr# spinel-bearing dunite. Orthopyroxenites formed at the expense of the pre-existing peridotite by reaction with hydrous orthopyroxene-saturated melts, which were produced by assimilation of dissolved pyroxene during the formation of the dunite. Refractory harzburgite, high-Cr# spinel-bearing dunite, and orthopyroxenite may have a genetic link to the late stage boninitic magmas in the crustal section of the EMO. In contrast, the Cpx-harzburgite was a residue related to mid-ocean ridge basalts (MORBs) or the “MORB-like” fore-arc basalt recently proposed by Reagan et al. (2010) from the Izu–Bonin–Mariana fore-arc. The mediumCr# spinel-bearing dunite can be caused by interaction with a melt transitional between MORB-like and boninitic melts. The lithological variations and their relationships in the upper mantle section of the EMO were caused by changes in magmatic compositions from mid-ocean ridge signatures to boninitic magmas, due to an increasing contribution of slab-derived fluids in an island arc setting. © 2010 Elsevier B.V. All rights reserved.

1. Introduction Ophiolites represent past fragments of oceanic lithosphere now tectonically emplaced on land. Field relationships combined with geochemical studies of volcanic sequences in ophiolites are commonly applied to reconstruct the magmatic and tectonic history of a part of the oceanic lithosphere (e.g. Dilek et al., 2008). A volcanic sequence in an ophiolite sometimes shows varying chemical characteristics, indicating spatial and/or temporal modification of the magmatic affinity in the ⁎ Corresponding author. Frontier Science Organization, Kanazawa University, Kanazawa 920-1192, Japan. Tel.: + 81 76 264 6513; fax: + 81 76 264 6545. E-mail address: [email protected] (T. Morishita). 0024-4937/$ – see front matter © 2010 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2010.10.003

ophiolite (e.g. Pearce et al., 1984). In contrast to the studies of volcanic rocks in the ophiolite, the evolutionary history of the mantle sections corresponding to changes in tectonic settings is not yet well understood. Many ophiolites are now proposed to form in the “supra-subduction zone” (SSZ) setting (Beccaluva et al., 1984, 1994; Miyashiro, 1973; Pearce et al., 1984; Shervais, 2001 and references therein). Shervais (2001) reviewed the petrological and geochemical signatures of SSZ ophiolites and suggested that SSZ ophiolites experienced a sequence of events during their evolution in response to the change in tectonic setting from oceanic lithosphere formed at mid-ocean ridges to the initiation of subduction as follows: birth, youth, maturity, death and resurrection. Studies of mantle sections from SSZ ophiolites provide important information on the maturing of mantle wedges above

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subduction zones, where continuous and/or episodic fluid fluxes are expected from the subducting oceanic plate. The Albanian ophiolites occur within the Dinaride–Hellenide segment of the Alpine orogenic system and represent the remnants of the Mesozoic Neo-Tethyan ocean (e.g., Shallo and Dilek, 2003; Dilek and Furnes, 2009). The Albanian ophiolites are generally divided into western- and eastern-types based on petrological and mineralogical data, as discussed by many authors (Bébien et al., 2000; Beccaluva et al., 1994; Bortolotti et al., 1996; Dilek et al., 2008; Hoeck et al., 2002; Nicolas et al., 1999; Shallo, 1990; Shallo et al., 1987, 1990 and references therein). The northern Albanian ophiolite belt, the Mirdita ophiolite, shows that MORB and SSZ affinities are dominant in the west and the east, respectively (Beccaluva et al., 1994; Bortolotti et al., 1996, 2002; Dilek et al., 2008; Shallo, 1990; Shallo et al., 1987, 1990). It should be emphasized that in this paper we specifically focus on the eastern region of the Mirdita ophiolite (the Eastern Mirdita Ophiolite). Beccaluva et al. (1994) pointed out that the mantle section of the eastern Mirdita ophiolite is characterized by strongly depleted signatures in melt components. Based on the petrology of the mantle section combined with geochemical signatures of volcanic rocks in the eastern part, they concluded that the eastern part of the Mirdita ophiolite was formed in supra-subduction settings. Bizimis et al. (2000) examined the trace element composition of clinopyroxene in some peridotites from the mantle sections of ophiolite complexes from the Hellenic Peninsula including one peridotite body (Bulqiza) from the Mirdita ophiolite. They suggested that the trace element composition of clinopyroxene in these ophiolites is similar to modern arc peridotites recovered from Izu-Bonin-Mariana arc (Parkinson and Pearce, 1998; Zanetti et al., 2006). Detailed geochemical mapping in several suprasubduction zone ophiolites has recently revealed the presence of

distinct units within the mantle section of an ophiolite (Arai et al., 2006; Batanova and Sobolev, 2000; Choi et al., 2008; Tamura and Arai, 2006; Uysal et al., 2009). We examine systematic petrological and mineralogical variations in the uppermost mantle section in the Eastern Mirdita ophiolite in the context of a maturing of mantle wedge.

2. Geological outline and sample descriptions The Mirdita ophiolite is located in the northern ophiolite belt of Albania (Fig. 1). Based on differences in the internal stratigraphy and chemical composition of the crustal unit, two types of ophiolites have been recognized in the Mirdita ophiolite, namely the Western Mirdita Ophiolite (WMO) and the Eastern Mirdita Ophiolite (EMO) (Beccaluva et al., 1994; Bortolotti et al., 1996; Dilek et al., 2008; Shallo, 1990; Shallo et al., 1987, 1990) Boninitic dikes and lavas crosscut and/or overlie earlier extrusive rocks in the EMO (Beccaluva et al., 1994; Dilek et al., 2008; Shallo et al., 1987). The crustal section of the WMO has MORB affinities, whereas that of the EMO shows predominantly SSZ geochemical affinities. The extrusive sequence in the EMO consists of pillowed to massive flows ranging in composition from basalt and basaltic andesite in the lower section to andesite, dacite, and rhyodacite in the upper part (Bortolotti et al., 1996; Dilek et al., 2008). Large peridotite massifs are exposed at the western and eastern ends of the Mirdita ophiolite. Plagioclase-bearing peridotites are frequently observed in the WMO, whereas harzburgite is dominant in the EMO (Beccaluva et al., 1994, 1998; Beqiraj et al., 2000; Hoxha and Boullier, 1995). In this paper, we focus on three peridotite massifs (Bulqiza, Lure and Kukësi massifs) located in the EMO (Fig. 1). All massifs have basically the same petrological and mineralogical characteristics. The

Fig. 1. Simplified lithological map of the studied area showing sample localities discussed in this study. Inset map shows the distribution of a part of the Tethyan ophiolites in the Balkan Peninsula, with the Mirdita ophiolite in red.

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Bulqiza massif has economically important high-Cr/Al chromitite ores (Beccaluva et al., 1998; Beqiraj et al., 2000). Systematic lithological variations in the mantle section with proximity to the crustal section have previously been recognized (Beccaluva et al., 1998; Beqiraj et al., 2000; Hoxha and Boullier, 1995). We also confirmed systematic lithological variations in the mantle section: clinopyroxene porphyroclast-bearing harzburgites (Cpxharzburgites hereafter, Fig. 2a) are sometimes observed in the eastern margin of massifs (01Blq, 02Lur), i.e. the basal part of the mantle section, whereas harzburgite and dunite are dominant in the upper parts of the mantle section (Beccaluva et al., 1998; Beqiraj et al., 2000; Dilek and Morishita, 2009; Hoxha and Boullier, 1995). Cpx-harzburgites have a porphyroclastic texture. Clinopyroxene occurs as both porphyroclastic grains and their recrystallized fine grains. The relationships between Cpx-harzburgite and other lithologies are not well observed in this study. The lithological boundary between dunites and harzburgites is usually sharp and is sometimes nearly parallel to the foliation plane defined by mineral orientations. Dunite also frequently occurs as small bodies with complicated irregular boundaries with harzburgites (Fig. 2b). Harzburgite shows granular to porphyroclastic textures. Chromitite layers a few cm thick are frequently observed in dunite and usually occur parallel to each other and to the lithological boundary between dunite and harzburgite (Fig. 2c). Chromitite layers are occasionally tightly folded in dunites (Fig. 2c). It is interesting to note that inclusions of silicate minerals, such as amphibole, orthopyroxene, clinopyroxene, and their secondary minerals (e.g., chlorite and serpentine), are commonly found within chromian spinels in harzburgites near dunite (Fig. 3a).

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Orthopyroxenite dikes/layers a few cm to 3 m wide are frequently observed in the uppermost section of the mantle sequence (Beccaluva et al., 1998; Dilek and Morishita, 2009) (Fig. 2d). They rarely occur as layers nearly parallel to the foliation and lithological boundaries in the host peridotites, and more frequently occur as dike-like features cutting all lithological boundaries at high angles (Fig. 2b), indicating that they are related to late melt migration through the mantle section. Fewer deformation textures are observed in orthopyroxenites. Orthopyroxenites mainly consist of coarse-grained orthopyroxene (up to 10 cm across) with small amounts of spinel and olivine. Olivines sometimes show resorbed textures in large orthopyroxene grains (Fig. 3b). Large orthopyroxenes have many clinopyroxene exsolution lamellae. Dark brown spinel is commonly included in large orthopyroxene grains. Orthopyroxenites locally contain amphibole and/or clinopyroxene. Clinopyroxene sometimes occurs as veins along orthopyroxenites. Amphibole occurs as an interstitial phase along the grain boundaries of orthopyroxene and also as poikilitic phases including orthopyroxene grains (Fig. 3c). 3. Mineral chemistry 3.1. Analytical methods Major-element compositions of minerals were analyzed using an electron probe micro-analyzer (JEOL JXA-8800 Superprobe) at Kanazawa University. The analyses were performed under an accelerating voltage of 15 kV and beam current of 20 nA, using a 3 μm diameter beam. Natural and synthetic mineral standards were employed for all minerals.

Fig. 2. (a) Polished surface of a clinopyroxene porphyroclast-bearing harzburgite (01 Blq). (b) Field relationships between dunite, harzburgites and orthopyroxenite (01 Lur). (c) Occurrence of chromitite layers in dunite (03 Kuk) Chromitite is sometimes tightly folded (arrow). (d) Orthopyroxenite network (yellow arrows) in harzburgites (04 Blq). C = chromitite, cpx = clinopyroxene, D = dunite, H or Harz = harzburgite.

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Fig. 3. (a) Back-scattered electron image of a spinel grain with silicate mineral inclusions in harzburgites near a dunite. (b) Resorbed olivine in a large orthopyroxene grain in an orthopyroxenite (04Blq). (c) Poikilitic amphibole (light green phase) in orthopyroxenite. amph = amphibole, ol = olivine, OPX = orthopyroxene, spl = spinel.

JEOL software using ZAF corrections was employed. Details of EPMA were described in Morishita et al. (2003). Representative major element compositions of minerals are shown in Table 1. Rare earth element (REE) and trace element (Li, Ti, Sr, Y, Zr, and Nb) compositions of minerals were determined by 193 nm ArF Excimer laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS) at Kanazawa University (Agilent 7500S equipped with MicroLas GeoLas Q-plus +; Ishida et al., 2004). Clinopyroxenes were analyzed by ablating 30–80 μm spot diameter at 5– 10 Hz, depending on the size and trace element abundances of analyzed minerals. The NIST SRM 612 was used as the primary calibration standard and was analyzed at the beginning of each batch of b8 unknowns, with a linear drift correction applied between each calibration. The element concentrations of NIST SRM 612 for the calibration are selected from the preferred values of Pearce et al. (1997). Data reduction was facilitated using 29Si and 42Ca as internal standard elements, based on Si and Ca contents obtained by EPMA following a protocol essentially identical to that outlined by Longerich et al. (1996). Details of the analytical method and data quality were described in Morishita et al. (2005a,b). Representative analyses in trace element compositions of minerals are shown in Table 2. 3.2. Olivine The forsterite and NiO contents of olivine increase from Cpxharzburgite (91.1 ± 0.1, 0.38± 0.02 wt.%) to dunite-associated chromitite layers (94–95, 0.46–0.53 wt.%) through harzburgites (91–92, 0.38– 0.4 wt.%) (Fig. 4). The forsterite and NiO contents of olivine in a dunite directly contacting with harzburgites (91.3 ± 0.1, 0.31 ± 0.04 wt.%) are identical to those in the host harzburgite (91.4 ± 0.2, 0.30 ± 0.03 wt.%). The forsterite and NiO contents of olivine in chromitite layers are higher than those in the host dunites (96, 0.5–0.58 wt.%), because of equilibration with chromian spinel at low temperature conditions. The NiO contents of olivines with ragged rims enclosed by large orthopyroxene grains in orthopyroxenites are likely to be slightly higher (up to 0.47 wt.%) than the host olivines (usually less than 0.4 wt.%) (Fig. 4). 3.3. Orthopyroxene The Mg# (= 100 Mg/(Mg + Fe) atomic ratio) of orthopyroxene increases from Cpx-harzburgite to harzburgite. The Mg-number of large orthopyroxene grains in orthopyroxenites is lower in a cpx-bearing layer (90) than others (92). The Al2O3 content of orthopyroxene usually decreases from core to rim (from 2.7 to 1.8 wt.% for Cpx-harzburgite) in both peridotites and orthopyroxenites. The Al2O3 contents of the cores of

orthopyroxene porphyroclasts decrease from Cpx-harzburgite (3 wt.%) to harzburgite (1.3 wt.%) (Fig. 5). The Al2O3 content of the cores of large orthopyroxene grains in orthopyroxenites without cpx and amphibole is usually 1.5 wt.%. Those surrounded by amphibole in amphibole-bearing orthopyroxenites reached 2 wt.%. The Al2O3 content of orthopyroxenite surrounded by amphibole in orthopyroxenites is higher (N2 wt.%) than others (1.5 wt.%). The TiO2 content of orthopyroxene is usually lower than the detection limit of the EPMA (b0.04 wt.%). Orthopyroxene inclusions within spinel in harzburgites are low in TiO2 (lower than the detection limit of the EPMA, b0.04 wt.%) and Al2O3 (0.3 wt.%) (Fig. 6a). 3.4. Clinopyroxene The Al2O3 content of clinopyroxene is 2 wt.% for Cpx-harzburgite, c.a. 1 wt.% for harzburgite and dunite, and 1.5–1.8 wt.% for cpx-bearing orthopyroxenite. The TiO2 and Na2O contents of clinopyroxene are usually lower than the detection limit (b0.04 wt.%) of the EPMA except for the TiO2 content of dunite and the Na2O content of harzburgite. Clinopyroxene in the cpx-bearing orthopyroxenite is slightly higher in Na2O (c.a. 0.15 wt.%) than in other samples (usually b0.05 wt.%). Clinopyroxene inclusions within spinel in harzburgites are low in TiO2 (b0.04 wt.%), Al2O3 (0.4 wt.%) and Na2O (b0.03 wt.%). Light REE (LREE) contents of clinopyroxene in the cpx porphyroclast-bearing harzburgite and one harzburgite are lower than the detection limit (0.01–0.02 ppm) of the analyses. Their chondrite (CH)-normalized REE patterns are, therefore, fractionated in LREEs (Fig. 7a). The CH-normalized REE patterns of clinopyroxene in other harzburgites are enriched in LREE compositions although middle REEs are lower than the detection limit (0.02 ppm) of the analyses. Their heavy REE (HREE) compositions are more depleted than those in the other harzburgites (Fig. 7b). Sr content is lower in the Cpx-harzburgite than others (Fig. 7b). Primitive mantlenormalized trace element patterns of clinopyroxene in dunites and harzburgites show a positive Ti anomaly (Fig. 7b). 3.5. Spinel It is noted that the Cr# [= Cr/(Cr+ Al) atomic ratio] of spinel with silicate inclusions in harzburgites is heterogeneous within each grain, as well as among different grains (0.4–0.7). Conversely, those in other samples are homogeneous. Despite the heterogeneities in some samples, the Cr# of spinel increases from Cpx-harzburgite (0.48) to dunite (0.78–0.82) through harzburgites (Fig. 8). The Cr-number of spinel in orthopyroxenite ranges from 0.5 to 0.7. The TiO2 content of spinel in clinopyroxene porphyroclast-bearing harzburgites,

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219

Table 1 Representative major element compositions of minerals. Sample

Lithology

Mineral

Anal#

SiO2

TiO2

Al2O3

Cr2O3

FeO*

MnO

MgO

CaO

Na2O

K2O

NiO

Total

Mg#

Cr#

01Blq

Cpx-harz

Olivine Std Spinel Std Opx core Opx rim Cpx core Olivine Std Spinel Std Olivine Std Spinel Std Cpx Olivine Std Spinel Std Olivine Std Spinel Std Olivine Std Spinel bright Spinel normal Spinel dark Amph inc Opx inc Cpx inc Cpx Cpx Olivine Std Spinel normal Spinel darc Amph inc Opx inc Cpx inc Cpx Opx core Opxc rim Olivine Std Spinel Std Olivine Std Spinel Std Opx core Opx rim Cpx Olivine-host Olivine-fine Spinel Std Opx core Opx rim Opx* Amphibole

N=7

41.0 0.2 b 0.03

b 0.04

b0.03

b0.1

b 0.03

40.5 1.7 0.70 0.29 0.71 b0.1

b 0.04

b0.03

b 0.03

0.68 0.32 23.9 b 0.04

b0.03 b0.03 0.03 b0.03

b 0.03 b 0.03 0.00 b 0.03

0.38 0.02 0.10 0.02 0.12 b 0.1 b 0.1 0.51 0.00

0.482 0.024 0.150 0.094 0.190 0.845

11.0 0.11 b0.03

58.7 0.35 b0.1

b 0.04

b0.03

b 0.03

b0.03

b 0.03

0.660 0.009 0.942

0.782 0.002 0.639

0.15 0.01 0.05 b 0.04

10.6 0.40 1.0 b0.03

57.2 0.58 0.59 b0.1

0.09 0.02 b 0.04

b0.03

b 0.03

25.5 b 0.04

0.05 b0.03

b 0.03 b 0.03

9.6 0.25 b0.03

61.1 0.48 b0.1

b 0.04

b0.03

b 0.03

b0.03

b 0.03

0.15 0.01 b 0.04

8.60 0.37 b0.03

60.4 0.90 b0.1

0.07 0.03 b 0.04

b0.03

b 0.03

b 0.04

b0.03

b 0.03

b 0.04 b 0.04 b 0.04 0.08 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04

15.5 18.2 23.7 5.6 0.53 0.41 1.2 1.2 b0.03

53.0 50.7 44.7 2.1 0.84 0.78 0.54 0.48 b0.1

b 0.04 b 0.04 b 0.04 12.6 0.26 24.4 23.7 23.8 b 0.04

b0.03 b0.03 b0.03 0.52 b0.03 b0.03 0.03 0.03 b0.03

b 0.03 b 0.03 b 0.03 0.20 b 0.03 b 0.03 b 0.03 b 0.03 b 0.03

b 0.04 b 0.04 0.07 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04

16.4 24.5 6.8 0.54 0.86 1.2 1.4 1.1 b0.03

52.5 44.0 2.1 0.77 0.94 0.62 0.57 0.41 b0.1

b 0.04 b 0.04 12.7 0.35 24.0 24.2 0.53 0.71 b 0.04

b0.03 b0.03 0.701 b0.03 b0.03 0.06 b0.03 b0.03 b0.03

b 0.03 b 0.03 0.23 b 0.03 b 0.03 b 0.03 b 0.03 b 0.03 b 0.03

0.573 0.023 0.967 0.962 0.003 0.670 0.011 0.950 0.002 0.554 0.027 0.913 0.001 0.503 0.524 0.551 0.947 0.925 0.952 0.941 0.947 0.919 0.001 0.542 0.610 0.943 0.934 0.950 0.950 0.923 0.922 0.913

0.784 0.007 0.272

0.15 0.01 b 0.04

0.682 0.547 0.171 0.489 0.424 0.251 0.215 0.197 0.684

0.04 0.02 b 0.04

16.6 0.28 b0.03

51.4 0.50 b0.1

b 0.04

b0.03

b 0.03

b 0.04

b0.03

b 0.03

0.495 0.017 0.914

0.675 0.005 0.774

0.05 0.03 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 b 0.04 0.10

21.3 1.3 1.3 0.98 1.1 0.00 0.00 18.0 3.0 1.2 0.48 2.2 11.0

47.6 1.5 0.53 0.25 0.59 0.00 0.00 52.2 3.1 0.19 0.09 0.42 1.4

b 0.04

b0.03

b 0.03

0.30 0.03 b 0.1

0.76 0.50 23.9 b 0.04 b 0.04 b 0.04

b0.03 b0.03 0.05 b0.03 b0.03 b0.03

b 0.03 b 0.03 b 0.03 b 0.03 b 0.03 b 0.03

0.04 0.02 0.07 0.34 0.46 b 0.1

b0.03 b0.03 b0.03 1.7

b 0.03 b 0.03 b 0.03 0.66

b 0.1 b 0.1 b 0.1 0.12

0.541 0.026 0.918 0.918 0.953 0.92 0.92 0.52 0.05 0.92 0.92 0.92 0.93

0.600 0.022 0.217 0.147 0.261

0.69 0.30 0.64 12.5

100.4 0.5 100.1 0.4 100.6 101.0 100.7 100.1 0.3 99.2 0.3 100.2 0.2 99.2 0.3 99.7 99.3 0.7 99.2 0.7 99.8 0.3 99.4 0.5 100.4 0.2 98.9 98.6 98.8 97.4 101.5 99.8 99.7 100.1 100.5 0.6 98.9 99.4 98.0 101.2 100.0 99.9 99.3 100.0 101.1 0.4 99.6 0.4 100.9 0.4 99.4 0.4 100.4 100.1 97.9 99.9 100.0 100.7 0.2 99.3 98.8 99.4 96.0

0.911 0.001 0.626 0.007 0.917 0.913 0.942 0.957

0.15 0.02 b 0.04

50.2 0.3 14.1 0.3 34.3 34.8 17.7 53.9 0.45 13.7 0.21 52.4 0.34 11.8 0.55 17.1 53.5 0.39 13.9 0.30 52.7 0.27 11.2 0.60 50.4 0.17 10.4 10.9 11.9 21.2 36.3 17.9 17.8 18.1 51.1 0.39 11.3 13.4 21.0 36.5 18.1 17.8 34.6 35.1 50.9 0.17 10.3 0.41 50.8 0.23 11.6 0.54 34.8 35.1 17.7 50.7 50.6 10.9 1.2 34.9 35.1 34.5 19.2

b0.03

29.2 1.7 2.7 1.8 2.0 b0.03

0.11 0.03 0.24 0.02 0.14 0.15 b 0.1 b 0.1

b 0.04

0.04 0.01 b 0.04 b 0.04 0.04 b 0.04

8.7 0.1 15.9 0.4 5.5 5.9 1.9 4.3 0.29 15.2 0.17 5.7 0.31 19.2 0.93 1.05 3.8 0.29 14.1 0.39 5.0 0.16 18.6 1.1 8.5 0.09 19.6 18.4 18.2 2.1 5.2 1.6 2.0 1.8 8.0 0.15 18.4 17.2 2.3 4.6 1.7 1.7 5.1 5.3 8.7 0.15 20.8 0.73 8.5 0.19 18.5 0.90 5.5 5.6 1.6 8.3 8.3 19.1 2.0 5.5 5.6 5.5 2.4

03Blq-L

Du/Chr (Chr)

Du/Chr (Du)

04Blq-F

Chromitite

Dunite layer

04Blq-B

03Blq-B

01Kuk-D

Harzburgite

Harzburgite

Dun/Harz (Dun)

Dun/Harz (Harz)

03Blq

Orthopyroxenite

N=6 21–7 24–7 42–13 N=4 N=8 N=6 N=6 74–7 N=7 N = 11 N = 10 N = 13 N=7 82–8 84–8 86–8 70–7 71–7 32–4 46–5 64–7 N = 16 45–4 49–4 87–10 102–11 103–11 31–2 41–3 26–3 N=4 N=6 N=9

56.4 57.6 54.2 41.4 0.09 b 0.03 41.5 0.16 b 0.03 54.4 41.4 0.31 b 0.03 41.4 0.16 b 0.03 41.0 0.05 b 0.03 b 0.03 b 0.03 52.8 58.2 54.6 54.3 54.6 40.9 0.20 b 0.03 b 0.03 52.0 58.3 54.3 54.4 57.0 57.2 41.0 0.24 b 0.03

N=8

41.1 0.26 b 0.03

100–6 102–6 103–9 80 179–18 N=6

57.3 57.5 52.9 40.5 40.4 b 0.03

21–4 24–4 16 133

56.6 57.0 55.8 46.9

0.28 0.03 0.10 0.01 0.32 0.04 b 0.1 0.05 0.02 0.29 0.02 b 0.1 0.35 0.03 0.12 0.02 0.39 0.29 0.26 0.05 0.17 b 0.1 b 0.1 b 0.1 0.12 0.02 0.27 0.29 b 0.1 0.14 b 0.1 b 0.1 0.12 0.14 0.16 0.05 0.33 0.03 0.14 0.05 0.31 0.06 0.13 0.14 b 0.1 0.15 0.15 0.3 0.05 0.13 0.13 0.10 0.02

0.46 0.01 b 0.1 b 0.1 0.58 0.03 b 0.1 0.53 0.01 b 0.1 0.38 0.02 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 0.40 0.03 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 b 0.1 0.31 0.04 b 0.1

0.811 0.004

0.825 0.008

0.696 0.652 0.558 0.201 0.516 0.558 0.233 0.207

0.66 0.05 0.10 0.12 0.11 0.08

FeO* is total iron. Mg# = Mg/(Mg + Fetotal) atomic ratio except for spinel (= Mg/(Mg+ Fe2+). Fe2+ in spinel was calculated from spinel stoichiometry). Cr# = Cr/(Cr+ Al) atomic ratio. Anal# = name of analytical point or analytical numbers (N) for average compositions. Harz = harzburgite, Du = dunite, Chr = chromitite, std = standard deviation, opx = orthopyroxene, Opx* = orthopyroxene included in amphibole, cpx = clinopyroxene. Spinel bright, normal, and dark = bright, normal and dark area within heterogeneous spinel grain, respectively; inc = inclusion within spinel; amph = amphibole.

harzburgites and orthopyroxenites is generally low (b0.05 wt.%) whereas that of dunite or chromitite is around 0.15 wt.%. The YCr [= 100 Cr/(Cr + Al+ Fe3+) atomic ratio] of spinel is slightly higher in dunites and chromitites (3–6) than in Cpx-harzburgites, harzburgites and orthopyroxenites (0.5–2.5). Spinels with silicate mineral inclusions in harzburgites are also heterogeneous in YCr (1–6).

3.6. Amphibole Following the classification of Leake et al. (1997), amphibole in the orthopyroxenites are edenite, edenitic hornblende, magnesio hornblende and tremolitic hornbrende. Amphibole inclusions within chromian spinel in harzburgites are magnesio hornblende, tremolitic

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6

Table 2 Representative trace element compositions of clinopyroxene. 01Blq Cpx-Harz 42-13

04Blq-B harzburgite 46-5

Li Ti Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

10 159 0.16 1.02 0.05 0.05 0.21 b0.01 b0.01 b0.01 b0.02 b0.02 b0.01 0.015 0.01 0.12 0.03 0.16 0.03 0.22 0.03

4 89 0.12 0.40 0.06 0.05 b 0.05 b 0.01 b 0.01 b 0.01 b 0.02 b 0.02 b 0.01 b 0.02 b 0.01 0.06 0.01 0.06 0.01 0.09 0.02

64-7

03Blq-B Harzburgite 31-2

03Blq-L Dunite 74-7

5 86 0.15 0.37 0.06 0.04 b0.05 b0.01 b0.01 b0.01 b0.02 b0.02 b0.01 b0.02 b0.01 0.04 0.01 0.06 0.01 0.10 0.02

4 52 0.25 0.18 0.07 0.07 0.20 0.29 0.01 b0.01 b0.02 b0.02 b0.01 b0.02 b0.01 0.01 0.01 0.04 0.01 0.08 0.02

0.61 296 0.86 0.56 0.08 0.06 0.03 0.08 b0.01 b0.01 b0.02 b0.02 b0.01 b0.02 b0.01 0.07 0.02 0.09 0.02 0.11 0.02

Anal# = name of analytical point (the same as Table 1). Harz = harzburgite.

hornblende and tremolite. The TiO2, Na2O, and K2O contents of amphiboles are usually low, b0.1 wt.%, b2 wt.% and b1 wt.%, respectively, and the K/(K+ Na) atomic ratio is b0.25. The Cr2O3 content is around 1.5 wt.%. Amphibole inclusions within spinel in harzburgite are low in Al2O3 (b7 wt.%), TiO2 (b0.1 wt.%) (Fig. 6b), Na2O (b1 wt.%) and K2O (b0.2 wt.%).

4. Discussion The petrology of the mantle section of the Eastern Mirdita ophiolite is characterized as follows (Dilek and Morishita, 2009). Structurally, cpx-harzburgites occur lower in the mantle section, while harzburgite and dunite are more abundant higher in the section. Dunite is associated with chromitite layers. Orthopyroxenite crosscuts all lithologies of dunite (chromitite)-harzburgite in the late stage. We discuss petrogenesis of (1) Cpx-harzburgites, (2) harzburgite-dunites (chromitites) and (3) orthopyroxenites. 0.6

0.5 alternating with chromitite layers

NiO wt. %

0.4

0.3

0.2

0.1

0 90

cpx-Harz 01-Blq 02-Lur

91

Harz

Dunite 03L-Blq

04B-Blq 03B-Blq

04F-Blq

01D-Kuk

01D-Kuk

92

93

94

Orthopyroxenite Fine grains Coarse-grains & host grains

95

96

Fo olivine Fig. 4. Relationships between Fo content and NiO wt.% of olivine in peridotites, orthopyroxenites and their host peridotites. Harz = harzburgite.

Al2O3 wt. %, Orthopyroxene

Sample # Lithology Anal #

cpx-Harz

SWIR

Harz 04B-Blq

01-Blq 02-Lur

5

03B-Blq 01D-KUK

MAR 1274A

4

CIR MAR OCC: Atlantis Massif

3

2

Garret IBM

1

0

Hess Deep

0

0.5

1

Cr# (Cr/(Cr+Al) atomic ratio, Spinel Fig. 5. Correlations between Al2O3 content (wt.%) of orthopyroxene porphyroclasts and the Cr# of spinels. Data source for the mid-ocean ridge systems: South Indian Ridge (SWIR), Atlantis II Fracture Zone, Dick (1989) and Seyler et al. (2003); Central Indian Ridge (CIR), Hellebrand et al. (2002), Morishita et al. (2009); Mid-Atlantic Ridge (MAR), Atlantis Massif (oceanic core complex; OCC), Tamura et al. (2008) and 1274A, Seyler et al. (2007); Garrett and Hess Deep, East Pacific Rise (EPR), Constantin (1999) and Dick and Natland (1996). Data source for the Izu–Bonin–Mariana forearc (IBM), Parkinson and Pearce (1998) and Zanetti et al. (2006).

4.1. Origin of Cpx-harzburgites in the lower part of the mantle section: a residue of less-fluxed partial melting related to mid-ocean ridge basalts or fore-arc basalts? Clinopyroxenes in Cpx-harzburgites are characteristically depleted in LREEs and other incompatible elements, indicating that they are a simple residue of partial melting and melt extraction from mantle materials rising adiabatically beneath ocean ridges (Hellebrand et al., 2002; Johnson et al., 1990). Based on the method of Hellebrand et al. (2001), the Cpx-harzburgites are the residue after 17% partial melting. This is consistent with the fractional melting model (Ozawa, 2001) of depleted MORB mantle (DMM: Workman and Hart, 2005) at spinelperidotite stability conditions (Fig. 9a). Cpx-harzburgite mineral compositions plot within the compositional range of depleted abyssal peridotites in mid-ocean ridge systems (Figs. 5, 7, 8), with the exception of the Mg# of clinopyroxene (Fig. 10). The Mg# of clinopyroxene is higher in the Cpx-harzburgite (and other samples) than in abyssal peridotites from mid-ocean ridge systems, when compared at the same Cr# of clinopyroxene and Fo content of olivine (Fig. 10b). The high-Mg# of clinopyroxene in the studied samples can be explained by a temperature-dependent Fe– Mg exchange reaction between olivine and clinopyroxene under subsolidus conditions (Kawasaki and Ito, 1994) (i.e., clinopyroxenes in the studied samples equilibrated at lower temperature than those in the abyssal peridotites from mid-ocean ridge systems) because of the differences in exhumation history near mid-ocean ridges and at convergent margins. We conclude that geochemical characteristics of the Cpx-harzburgite in the EMO are very similar to those of depleted abyssal peridotites in mid-ocean ridge systems. It is, however, not necessary that the Cpx-harzburgite formed beneath a mid-ocean ridge; Reagan et al. (2010) recently reported MORB-like basaltic magmatism followed by boninitic magmatism in the Izu–Bonin–Mariana fore-arc regions. They postulated that the MORBlike tholeiitic basalts were the first lavas to erupt after the oceanic plate began to subduct and termed them “fore-arc basalt (FAB)”. Stern and Bloomer (1992) predicted a localized extensional zone directly above the shallow dipping slab in a compressional setting. This conceptual model was examined with visco-elastoplastic models by Hall et al. (2003) and a following paper (Gurnis et al., 2004). According to this model, adiabatic decompressional melting for the MORB-like melts

T. Morishita et al. / Lithos 124 (2011) 215–226

10

a: Orthopyroxene

a cpx-Harz (01-Blq) Harz (04B-Blq)

Chromitite (Hess Deep) Chondrite-normalized

0.2

Troctolite (Hess Deep) TiO2 wt. %

221

Harz (Atlantis Massif) 0.1

Dunite (03L-Blq)

Harz (03B-Blq) 1

0.1

This study Detection limit 0.01 0

0

0.5

1

1.5

2

La Ce Pr Nd

Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

2.5 1

Al2O3 wt. %

b

b: Amphibole

Primitive mantle-normalized

5

Troctolite (Mid-Cayman) Chromitite (Mid-Cayman)

4

TiO2 wt. %

Troctolite (Hess Deep) 3

Chromitite (Hess Deep) 2

0.1

0.01

cpx-Harz (01-Blq) Harz (04B-Blq) Harz (03B-Blq) Dunite (03L-Blq)

1

This study 04B-Blq 03B-Blq

0.001

Harz (Atlantis Massif)

0 0

5

10

15

Nb La Ce Sr Nd Zr Sm Eu Gd Ti Dy Y

Er Yb Lu

Fig. 7. Chondrite-normalized REE patterns (a) and primitive mantle-normalized trace element patterns (b) of clinopyroxenes in cpx porphyroclast-bearing harzburgites (cpx-Harz), harzburgites (Harz) and dunite (Dunite).

Al2O3 wt. % Fig. 6. Compositional relationship between TiO2 wt.% and Al2O3 wt.% of orthopyroxene and amphibole inclusions within spinel in harzburgites from this study and the Atlantis Massif of the Mid-Atlantic Ridge (Tamura et al., 2008). Those in chromitites and associated troctolite from ocean floor (Arai and Matsukage, 1998) are also shown. Harz = harzburgite.

occurred due to infiltration of hot mantle into the extensional region. Geochemically, no significant differences in major and trace element compositions of residues after partial melting are expected between mid-ocean ridges and the extensional zone at the fore-arc setting. In fact, Reagan et al. (2010) pointed out that some ophiolites, including the Mirdita amphibolites, were originated by near-trench volcanism at subduction initiation. We cannot say definitively whether the Cpxharzburgite is a residue related to MORB or FAB, and further studies are required as more generally discussed later. In the next section, we also discuss a possibility that some dunites (chromitites) in this study were related to the FAB. In addition to the discussion above, it is worthwhile to compare the EMO and WMO mantle sections because MORB-like igneous rocks dominate in the crustal section of the WMO, although the relationship between EMO and WMO has been a matter of debate (Bébien et al., 1998, 2000; Bortolotti et al., 2002; Hoeck et al., 2002; Meshi et al., 2010; Nicolas et al., 1999; Shallo and Dilek, 2003; Tremblay et al., 2009). In the Shebenik massif, located south of the Bulqiza massif (Fig. 1), Bébien et al. (1998) suggested that plagioclase wehrlites, troctolites and gabbros above highly depleted harzburgites and dunites were products of the

crystallization of MORB affinity melts. Hoeck et al. (2002) proposed a systematic variation in petrography and geochemistry from north to south in the western belt, with increasingly SSZ signatures towards the south, based on studies in the western-type of the southern Albanian ophiolites (Voskopoja ophiolites) combined with results from the Pindos ophiolites. On the other hand, Nicolas et al. (1999) and Meshi et al. (2010) suggested that the differences between the WMO and EMO were due to successive episodes of magmatic and amagmatic extension in a slow-spreading mid-ocean ridge environment. Further studies on petrology and mineralogy in the mantle section of the WMO can provide useful lines of information about the variation of the mantle section with respect to differences in the crustal section. 4.2. Origin of refractory harzburgites and dunites including chromitites: implications for influx melting beneath island arc setting The presence of refractory peridotites consisting of harzburgites and dunites (including chromitites) was previously recognized as a dominant characteristic of the uppermost section in the EMO (Beccaluva et al., 1994, 1998; Bizimis et al., 2000). Mineral compositions in harzburgites and dunites (+ chromitites) are generally outside of the compositional range of those in abyssal peridotites from mid-ocean ridge systems, but are similar to those of fore-arc peridotites (Figs. 5, 8 and 10). Clinopyroxenes from harzburgites and dunites have lower heavy REE concentrations and extremely low middle REE concentrations (usually lower than the detection limit of the analyses), whereas LREE (+ Sr) are

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100

a

90 80

10

Boninite-IBM & Others

1st stage: less flux

Boninite Dunite Fore-arc Chr peridotites

80

Opxnite Harz

70

Dunite (01D-Kuk)

60

Cr# 50

Fore-arc

High-Cr du. (this study) Med.-Cr du. (this study)

Cr# 50 cpx-Harz

40

40

30

30

IBM Dunite

Abyssal peridotites

20

20

10

10

5%

1 10%

15% (2nd stage)

0.1

17%

0.01

Abyssal

Dunite (03L-Blq)

0.001

Mg#

0

0 100

50

0

10

Mg#

Fig. 8. Compositional relationship between Cr# (= Cr/(Cr + Al) atomic ratio) and Mg# (= 100Mg/(Mg + Fe2+) atomic ratio) of spinel. (a) Compositional variation of Cpxharzburgite (cpx-Harz), harzburgite (Harz), dunite (Dunite), chromitite (Chr), and orthopyroxenite (Opxnite) in the mantle section of the EMO. Compositional range of a boninitic dike from the Mirdita ophiolite is also shown (Boninite) (unpublished data). (b) Comparison between the studied dunite, the Izu–Bonin–Mariana (IBM) dunite, and boninites from the IBM and other localities. Data for abyssal peridotites from Dick (1989), Arai and Matsukage (1998), plagioclase-poor samples of Constantin (1999), Hellebrand et al. (2002), harzburgites of Seyler et al. (2003); for fore-arc peridotites from Ishii et al. (1992) and Parkinson and Pearce (1998); and for boninites from Kuroda et al. (1978), Walker and Cameron (1983), Bloomer and Hawkins (1987), Falloon et al. (1989), van der Laan et al. (1992) and Sobolev and Danyushevsky (1994). The IBM dunite of Ishii et al. (1992) in (b) was averaged data.

enriched in some samples, as already suggested by Bizimis et al. (2000). HREE concentrations are most affected at high-degree of partial melting, whereas LREEs are likely to be susceptible to changes during/after melting events. Chromian spinels in some harzburgites contain hydrous silicate inclusions. The presence of hydrous silicate inclusions in the spinel does not directly suggest the infiltration of the hydrous fluids/melts since hydrous silicate inclusions in chromian spinel from “harzburgites” were also reported from the Atlantis Massif along the Mid-Atlantic Ridge where reactions between peridotites and highly evolved melts from MORB compositions were expected (Tamura et al., 2008). However, gabbroic rocks do not always occur close to the harzburgites in the EMO. Hydrous silicate inclusions within spinels were also reported in oceanic chromitite–dunite (+ troctolites) (Arai et al., 1997). Arai et al. (1997) suggested that the hydrous silicate inclusions crystallized from highly evolved melts after melt-peridotite interactions coupled with zonerefining effects (Harris, 1957; Kushiro, 1968) in a stagnant or failed melt conduit. Although chemical compositions of hydrous silicate inclusions within spinel do not directly reflect a parent melt composition which reacted with the host peridotites to result in the formation of dunite, occurrences and main mineral assemblages are similar between the mid-ocean ridge samples and the studied samples. Irrespective of the similarities in occurrences of silicate inclusions and their host rocks, chemical characteristics of silicate inclusions, such as TiO2 content, are apparently distinctive from those in the mid-ocean ridge samples (Fig. 6). These chemical and petrological features in harzburgites can be produced as a result of enhanced partial melting of depleted peridotites caused by the infiltration of a hydrous LREE-enriched, TiO2-poor flux, which is different from MORB-related melts. We initially applied an open system melting model using the spreadsheet of Ozawa (2001) in order to reveal chemical signatures of influx fluids/melts. We assumed that the starting composition was the residue after 15% fractional melting from DMM and that the degree of

Ce

Sr

Nd

Zr

Sm Eu

Ti

Dy

Y

Er

Yb

2nd stage: infulx melting Influx total melt segregated from the system (melting model)

Primitive mantoe-normalized

50

cpx-Harz (01-Blq) Harz (04B-Blq) Harz (03B-Blq)

20%

P & P 98 Ishii et al (averaged)

La 0 100

Start

90

70 60

b

Primitive mantoe-normalized

100

1

Calc-melt (dunite-cpx)

Start

0.1

dunite cpx

0.01 Influx Calc-cpx (melting model)

0.001 La

Ce

Sr

Nd

Zr

Sm Eu

Ti

Dy

Y

Er

Yb

Fig. 9. Observed and modeled primitive mantle-normalized trace element patterns for clinopyroxene. Initial bulk compositions and initial mineral mode for the 1st stage are from Workman and Hart (2005), while those for the 2nd stage are the result after 15% fractional melting of the 1st stage. Melting stoichiometries for the 1st and 2nd stages at spinelperidotite conditions are olivine 0.05, spinel 0.05, orthopyroxene 0.3, clinopyroxene 0.6; and olivine — 0.1, spinel 0.1, orthopyroxene 0.4, clinopyroxene 0.6, respectively. Mineral/melt partition coefficients are listed by Ozawa and Shimizu (1995) and Ozawa (2001). The calculated melt compositions equilibrated with clinopyroxene (calc-melt, dunite-cpx) and total melt segregated from the system using this model are also shown in (b). Influx composition for La, Ce, Sr and Zr was optimized as follows: critical melt fraction α= 0.01; influx rate β= 1; melt separation rate after trapped melt τ= 0; trapped melt crystallization reaction stoichiometry = olivine 0.4, spinel 0.05, orthopyroxene 0.2, clinopyroxene 0.35. Influx compositions for HREEs+ Y and Ti were assumed to be 0.01 times and 0.005 times the primitive mantle values, respectively.

partial melting was 5% for the 2nd stage of melting (Fig. 9b). Other parameters are shown in the caption of Fig. 8. We optimized influx fluid/melt compositions for La, Ce, Sr and Zr under a critical open system melting to fit their contents to observed clinopyroxene in harzburgite (Fig. 8b), whereas HREE and Ti contents of the influx fluids/melts were simply assumed to be low, 0.01 and 0.005 times the primitive mantle values of McDonough and Sun (1995), respectively. An estimated influx fluid/melt has high LREE/HREE ratios as expected. Regardless of our preliminary estimation, total melt compositions from the systems using the model show some similarities to the calculated melt equilibrated with clinopyroxene in dunite, such as high LREE/HREE ratio and positive Ti anomaly, but large differences in Sr content. Both calculated melts and simulated melts are different from volcanic rocks in the crustal section of the EMO, probably indicating that magma compositions changed during transport to the upper crustal level as a result of crystallization and/or interactions with lower-crustal materials. Since compositions of influx fluids/melts

T. Morishita et al. / Lithos 124 (2011) 215–226

0.4

Zhou, 1994). It is interesting to note that there are large differences in the Fo content of olivine and the Cr# of spinel in dunite samples: one group is characterized by high-Cr# spinel and high-Fo olivine, and the other by medium-Cr# spinel and low-Fo olivine (Figs. 4 and 8). Two distinctive melts, at least, were required for the formation of these dunites. The Cr# of spinel in a boninitic dike in the EMO is the same as that in the high-Cr# spinel-bearing dunite (Fig. 8). One melt responsible for the high-Cr# spinel-bearing dunite (and refractory harzburgite as well) can, therefore, be assumed for magmas related to boninite activities that were recorded in the crustal section of the EMO (Dilek et al., 2008). On the other hand, the Cr# of spinel in the other group (Cr# 0.67) is higher than those in typical abyssal peridotites in the mid-ocean ridge systems, and is generally lower than the majority of those in boninitic magmas (Fig. 8b). Reagan et al. (2010) reported a transition magma between the MORB-like FAB and boninitic compositions in the IBM fore-arc. The medium-Cr# spinel-bearing dunite could be a melt conduit for transition magmas. The wide range of variation in the Cr# of spinel in dunite from the IBM (Fig. 8b) was probably caused by changing melt compositions in the fore-arc setting throughout this time.

Cr/(Cr+Al) atomic ration, Clinopyroxene

a 0.3

MAR 1274 MAR Atlantis M Hess Deep Garret

0.2

0.1

0.0 0.88

IBM SWIR

CIR

0.90

Harz

cpx-Harz 01-Blq 02-Lur 0.92

0.94

04B-Blq 01D-KUK 03B-Blq 0.96

0.98

4.3. Origin of the orthopyroxenites: Silica-enrichment in the uppermost mantle section beneath island arc setting

Mg/(Mg+Fe) atomic ratio, Clinopyroxene 0.94

cpx-Harz

b

Mg/(Mg+Fe) atomic ratio, Olivine

01-Blq 02-Lur 0.93

Harz

04B-Blq 01D-KUK 03B-Blq 0.92

IBM

MAR 1274 1:1

0.91

Hess Deep 0.9 0.9

0.92

0.94

0.96

223

0.98

Mg/(Mg+Fe) atomic ratio, Clinopyroxene Fig. 10. Compositional relationships between Mg# and Cr# in clinopyroxene (a), and Mg# of coexisting clinopyroxene and olivine (b). Only data with Mg# N 0.9 and Cr# b 0.4 are shown. Data sources are the same as Fig. 8.

are also expected to change during interactions with the host rocks (Navon and Stolper, 1987), we need more systematic sampling from several localities including crustal sections to reveal the quantitative evolution of melt compositions though melt-rock interactions. Some dunites, particularly those associated with chromitites, have high-Fo and NiO contents of olivine, and high-Cr# of spinel. These characteristics might be explained if they are the residues after extremely high-degrees of partial melting (Kubo, 2002). However, the dunite with low-NiO and low-Fo olivine (Figs. 4 and 8) and the fact that the boundaries between dunite and harzburgite are usually very sharp suggest an origin as reaction products rather than residues after extremely low degree partial melting. Partial melts formed at high pressure in the mantle are saturated only in olivine at lower pressure, and will dissolve pyroxenes during melt/rock reaction resulting in dunite (Kelemen, 1990; Kelemen et al., 1995). Chromitites in upper mantle peridotites are generally interpreted to be formed by melt– mantle interactions (Arai, 1995, 1997a,b; Arai and Yurimoto, 1994;

Small ragged olivine grains in large orthopyroxene grains in orthopyroxenites are higher in NiO than other olivine grains in the host peridotites. Early crystallization of orthopyroxene implies that the melts for the formation of orthopyroxenites were saturated in orthopyroxene components, i.e., high-SiO2 content. This texture combined with the presence of amphibole indicates that the orthopyroxenites formed at the expense of the pre-existing olivine by reaction with hydrous orthopyroxene-saturated melts, resulting in an increasing NiO content in the residual olivine as the volume of the olivine was decreased (e.g., Kelemen et al., 1998). Assimilation of dissolved pyroxene for the formation of the dunite in the studied area, moving it towards opx saturation, may be a major mechanism for the formation of melts saturated in orthopyroxene, and could be the source for the orthopyroxenites (Kelemen, 1990). Furthermore, the effect of H2O substantially increases the maximum SiO2 content of olivine-saturated liquids. This is consistent with the orthopyroxenites being products of hydrous partial melting in the EMO uppermost mantle sections. This combined with the late formation and high-Cr# spinel (Fig. 8) indicates a genetic link to the late stage boninitic magmas in the crustal section of the EMO. If this is true, it is noted that geochemical signatures of boninitic melts may have been acquired during the interaction with surrounding depleted peridotites during the migration of melts through the mantle (Varfalvy et al., 1997). Further systematic studies on orthopyroxenites are needed to understand the origin of high-Mg andesite, including boninitic magmas. Progressive CaO enrichment due to the crystallization of orthopyroxene (and spinel) from orthopyroxene-saturated parent melts lead to the crystallization of clinopyroxene, and amphibolebearing orthopyroxenites were finally formed from the most evolved members of the melts related to the orthopyroxenites. Orthopyroxenites were commonly observed in the peridotite section in many ophiolites (e.g., Miyamori ophiolite of Japan, Ozawa, 1988; Leka ophiolite of the Norwegian Caledonides, Furnes et al., 1992; Maaløe, 2005; Bay of Island ophiolite of Newdoundland, Varfalvy et al., 1996, 1997; Josephine peirdotite of dismembered ophiolite in the Klamath Mountains, Kelemen and Dick, 1995; Oman ophiolite, Tamura and Arai, 2006; Coast Range ophiolite of California, Choi et al., 2008). Tamura and Arai (2006) examined orthopyroxenites with their related harzburgites and dunites from the northern part of the Oman ophiolite. They revealed that their suites are characterized by high-Cr# spinel and clinopyroxene with low abundances of middle-heavy REEs and LREE-enrichments, as compared with other samples, suggesting that the calculated melts from clinopyroxene compositions are similar to boninitic melts. This is consistent with the fact that boninitic magmas and their relatives were

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recently reported in the crustal sequence of the Oman ophiolite (Ishikawa et al., 2002; Yamasaki et al., 2006). Silica-enrichment of mantle peridotites, i.e., the formation of orthopyroxene, has also been reported in several sub-arc xenoliths (Arai and Kida, 2000; Arai et al., 2003, 2004; Ishimaru et al., 2007; McInnes et al., 2001). We would like to emphasize that silicaenrichment in the uppermost mantle section under island arcs is a ubiquitous phenomena caused by infiltration of silica-rich hydrous fluids/melts derived from subducting slabs. Orthopyroxene in some sub-arc xenoliths is, however, characterized by extremely low Al2O3, Cr2O3 and CaO contents, and is thought to have a metasomatic origin related to interaction with slab-derived fluids. Silica-enrichment observed in the uppermost section of the peridotite massifs in the EMO as well as other ophiolites may be applied more generally to speculate on subduction zone magmatism rather than simple metasomatic reactions due to infiltration of slab-derived fluids. 5. Concluding remarks This study revealed the presence of geochemically distinct peridotites in the mantle section of the EMO. One is the Cpx-harzburgite which is a residue of less-flux melting. The other is highly refractory harzburgite– dunite (+ chromitite) which was mainly formed by a high-degree of partial melting with LREE/HREE flux components, most likely related to boninitic magmatism in the mantle wedge. The latter is more common in the upper section of the EMO peridotite massifs. Orthopyroxenite is the later product of melt-rock reactions. These lithological variations and their relationships observed in peridotite massifs from the EMO can be explained by changes in magmatic compositions from mid-ocean ridge basalt signatures to arc-related magmas including boninitic magmas. The contributions of arc-related components increased, modifying the preexisting “mid-ocean ridge-like” mantles, and resulting in the formation of harzburgite–dunite (+ chromitite)–orthopyroxenite suites. It is not yet clear whether the pre-existing mantle is of true mid-ocean ridge origin or FAB-related. This scenario is basically consistent with the one that was speculated from volcanic sequences in the EMO (e.g. Dilek et al., 2008; Reagan et al., 2010). Detailed geological and geochemical studies in several ophiolites have revealed the presence of distinct peridotite units within the mantle section of ophiolites (Ahmed and Arai, 2002; Aldanmaz et al., 2009; Arai et al., 2006; Batanova and Sobolev, 2000; Caran et al., 2010; Choi et al., 2008; Ghazi et al., 2010; Melcher et al., 1997; Saccani et al., 2010; Tamura and Arai, 2006; Uysal et al., 2009). In the case of the Troodos ophiolite, which contains arc-signatured magmatic rocks in the crustal section (Miyashiro, 1973; Pearce and Robinson, 2010; Robinson and Malpas, 1990), Batanova and Sobolev (2000) reported that abyssal-type peridotites were distributed in the lower part of the peridotite body whereas highly refractory harzburgites and dunite are dominant in the upper part of the body. These lithological variations and their relationships in ophiolites can be explained by a shift in tectonic setting from mid-ocean ridge to island arc. If, however, the “MORB-like” FAB is a ubiquitous phenomenon at the initiation of subduction, we should reconsider our interpretation of ophiolite mantle sections. FAB magmatism at the earliest stage of subduction initiation is a new concept. More careful investigations of fore-arc peridotites and the mantle section of ophiolites are required to determine the differences between MORB and FAB, and whether FAB magmatism at subduction initiation is present at other locations. The frequency of highly refractory harzburgite (+ dunite, chromitite) and orthopyroxenite dikes in the mantle section can be a good indicator of arc-related magmatic modifications in the mantle, and is expected to vary from locality to locality. For instance, these lithologies are more common in the EMO than in the Oman ophiolite. This combined with the differences in crustal sections between ophiolites may reflect differences in the stage of life of the subduction zone: i.e., the Oman and the EMO may be snapshots of “baby” and “teenager” stages, respectively. The

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