Orthopyroxene-rich Rocks from the Sanbagawa Belt (SW Japan

Orthopyroxene-rich Rocks from the Sanbagawa

Belt (SW Japan): Fluid–Rock Interaction in the

Forearc Slab–Mantle Wedge Interface

Shunsuke Endo1*, Tomoyuki Mizukami2, Simon R. Wallis3,

Akihiro Tamura2 and Shoji Arai2

1Institute of Geology and Geoinformation, National Institute of Advanced Industrial Science and Technology

(AIST), Central 7, Tsukuba 305-8567, Japan, 2Earth Science Course, School of Natural System, College of Science

and Engineering, Kanazawa University, Kanazawa 920-1192, Japan and 3Department of Earth and Planetary

Sciences, Graduate School of Environmental Studies, Nagoya University, Nagoya 464-8602, Japan

*Corresponding author. Telephone: 81-29-861-2609. E-mail: [email protected]

Received November 29, 2013; Accepted May 27, 2015

ABSTRACT

The Western Iratsu body of the Sanbagawa belt (SW Japan) is a mafic–ultramafic complex that

underwent an initial metamorphism in the amphibolite facies and a subsequent metamorphism inthe eclogite facies, and represents a fossil forearc slab–mantle wedge interface in a developing

subduction zone. Two generations of orthopyroxene (Opx1 and Opx2) that were formed during the

amphibolite-facies (antigorite unstable) and eclogite-facies (antigorite stable) stages can be recog-

nized in the ultramafic domain. Opx1-rich rocks contain Ni-rich relict olivine (up to 0�7 wt % NiO)

and grade into dunite, suggesting that they represent metasomatic rocks derived from dunite.

Opx1 can be subdivided into two types: one (Opx1L) constitutes replacive harzburgite to orthopyr-

oxenite layers and the other (Opx1V) occurs in metasomatic reaction veins in dunite. Relativelyhigh formation temperatures (�750�C) of Opx1L imply that the relevant metasomatism in the ultra-

mafic domain took place before the juxtaposition with the mafic domain preserved in the Western

Iratsu body. Textural relationships and mineral trace element data suggest that Opx1L-rich rocks

were formed by reactive porous infiltration of a slab-derived hydrous melt or solute-rich fluid into

dunite. Subsequently, Opx1V-rich veins were formed by a prolonged flux of a Si-rich aqueous fluid

(sourced from the mafic domain) through brittle fractures in dunite during the amphibolite-faciesmetamorphism (�660�C and 1�2 GPa). The initial formation of Opx1V-chlorite-rich selvages along

the fluid conduits is likely to have limited the reaction between a Si-rich crustal fluid and host dun-

ite, and this process can be important during the early transportation of slab-derived components

into the mantle wedge. Lastly, Opx1L crystals locally show a textural replacement by Opx2 together

with antigorite, indicating recrystallization in the eclogite facies (�620�C and 1�6–1�8 GPa). The

Opx2-forming reaction is mainly localized in ductile shear zones, which correspond to major fluid

pathways in the partially serpentinized forearc mantle.

Key words: fluid; metasomatism; orthopyroxene; Sanbagawa belt; subduction zone

INTRODUCTION

The slab–mantle wedge interface is a site of intensive

chemical–mechanical interactions between mantle andcrustal rocks in the presence of a slab-derived hydrous

agent (fluids, melts or supercritical fluids). Silica meta-

somatism is one of the most important processes in the

mantle wedge immediately above the subducting slab,

typically resulting in the formation of talc- or orthopyr-

oxene-rich rocks (e.g. Manning, 1995). Talc- or orthopyr-

oxene-rich metasomatic rocks are thought to playcritical roles in subduction interface processes such as

mechanical coupling (Peacock & Hyndman, 1999;

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J O U R N A L O F

P E T R O L O G Y

Journal of Petrology, 2015, Vol. 56, No. 6, 1113–1137

doi: 10.1093/petrology/egv031

Original Article

Downloaded from https://academic.oup.com/petrology/article-abstract/56/6/1113/1526973by gueston 10 April 2018

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Moore & Lockner, 2007; Hirauchi et al., 2012) and recy-

cling of volatile and incompatible elements (Bebout &

Barton, 2002; Malaspina et al., 2006, 2009; Spandler

et al., 2008; Marschall & Schumacher, 2012).

Two scenarios can be considered for the formationprocess of orthopyroxene-rich rocks at the deep forearc

to subarc slab–mantle wedge interface: (1) silica meta-

somatism of peridotites [olivineþSiO2 (in melt or fluid)

! orthopyroxene]; (2) dehydration of silica-enriched

serpentinite (talcþ antigorite ! orthopyroxeneþH2O).

Both are likely to occur in nature, but natural occur-

rences of orthopyroxene-rich metasomatic rocksformed at the slab–mantle wedge interface are very

rare. This rarity may be due to the lack or scarcity of ex-

humation processes that favour their preservation.

However, examples of comparable rock types have

been described from (1) ultramafic massifs in continen-

tal subduction-type high-pressure (HP) and ultrahigh-pressure (UHP) orogens (Malaspina et al., 2006, 2009;

Scambelluri et al., 2006; Marocchi et al., 2007; Vrijmoed

et al., 2013), (2) supra-subduction zone ophiolites

(Sorensen & Barton, 1987; Berly et al., 2006) and (3)

xenoliths entrained in diatremes (Smith et al., 1999;

Facer et al., 2009). Orthopyroxene-rich rocks formed byhigh-pressure dehydration of serpentinite have also

been intensively studied in Cerro del Almirez, Spain

(Trommsdorff et al., 1998; Scambelluri et al., 2001;

Padron-Navarta et al., 2011), although this example is

believed to be a subducted oceanic serpentinite rather

than mantle wedge material (Marchesi et al., 2013).

Orthopyroxene-rich veins in mantle xenoliths en-trained in arc magmas have also been intensively

studied to understand fluid/melt–rock interactions in the

mantle wedge far above the slab surface. The proposed

origin of the vein-forming agents includes the follow-

ing: (1) slab-derived Si-rich hydrous melts/aqueous flu-

ids (Gregoire et al., 2001; McInnes et al., 2001; Franzet al., 2002; Arai et al., 2003; Ishimaru et al., 2007; Arai &

Ishimaru, 2008), (2) mantle-derived boninite-like melts

(Halama et al., 2009; Benard & Ionov, 2013); (3) host an-

desitic magmas (Ishimaru et al., 2007; Benard & Ionov,

2013). Multiple origins of fluids may complicate petro-

genetic interpretations of various types of orthopyrox-

ene-rich veins in mantle xenoliths.To understand early transportation processes of

slab-derived components into the mantle wedge, it is

important to shed light on the nature of hydrous meta-

somatism at the slab–mantle wedge interface preserved

in oceanic subduction-type orogens. The Sanbagawa

belt of SW Japan is one of the best-studied subduction-type orogens, but there have been few studies of this

belt that focus on this issue. Field geological mapping

of the Sanbagawa belt in central Shikoku shows that

ultramafic blocks occur exclusively in the higher pres-

sure region (from 0�8 to �3�0 GPa), suggesting that (1)

the surface of the original subducted slab was almost, if

not entirely, free from ultramafic rocks (typical of oce-anic lithosphere created at fast-spreading centers), and

(2) ultramafic rocks in the Sanbagawa belt represent a

suite of rock bodies that sample the entire depth range

(�30 to 100 km) of the forearc mantle where it formed

the hanging wall to the subduction zone (Aoya et al.,

2013a). Within the high-pressure region, the Western

Iratsu body is a unique mafic–ultramafic complex thatpreserves records of early and late periods of subduc-

tion zone development (Endo et al., 2009, 2012) and

thus provides an important opportunity to study the

slab–mantle interactions in an evolving subduction

zone environment. In this study we document the

discovery of various types of orthopyroxene-rich meta-

somatic rocks from the Western Iratsu body.The present study aims (1) to describe the field rela-

tions and petrological characteristics of the Western

Iratsu orthopyroxene-rich rocks, (2) to unravel the petro-

geneses of these rocks, and (3) to combine the newly

obtained information with the above-cited studies and

previous data on the Sanbagawa belt to further ourunderstanding of subduction interface processes, par-

ticularly concerning fluid–rock interaction.

GEOLOGICAL OUTLINE

Sanbagawa beltThe Sanbagawa belt of SW Japan represents a regionof Cretaceous high-P metamorphism resulting from the

subduction of the Izanagi oceanic plate beneath eastern

Asia (e.g. Wallis et al., 2009). It extends for more than

800 km along strike, and the maximum width reaches

�30 km in central Shikoku, where the belt is divided into

three tectonic units: the lowermost Oboke Unit, theBesshi Unit at intermediate structural levels, and the

uppermost Eclogite Unit (Fig. 1a and b). In the Besshi

and Eclogite Units, mineral assemblages in pelitic schist

have been used to define four post-eclogitic meta-

morphic zones: the chlorite, garnet, albite–biotite and

oligoclase–biotite zones, in ascending order of meta-

morphic grade (Higashino, 1990). The lithological andregional metamorphic sequences are locally overturned

by kilometer-scale recumbent folds (Banno et al., 1978;

Mori & Wallis, 2010) that developed during the main

phase of ductile deformation, referred to as DS (Wallis,

1990).

The main rock types in the Besshi and Eclogite Unitsare pelitic, mafic and siliceous schists, whose protoliths

are trench-fill sediments, altered mid-ocean ridge basalt

(MORB) and chert (Terabayashi et al., 2005; Aoya et al.,

2013a, 2013b). In addition, the presence of volumetric-

ally minor but widespread decimeter- to kilometer-scale

ultramafic blocks exclusively in the higher-grade (above

the garnet zone) regions (Fig. 1a) suggests tectonic en-trainment of hanging-wall mantle material into the

upper part of the subducted metasedimentary se-

quences (Maekawa et al., 2004; Aoya et al., 2013a).

Most small ultramafic blocks are serpentinite with meta-

morphic olivine (Kunugiza et al., 1986), but pristine peri-

dotite bodies (dunite–wehrlite suites) also occur in theEclogite Unit (e.g. Kunugiza et al., 1986; Hattori et al.,

2010). These peridotite bodies are associated with

1114 Journal of Petrology, 2015, Vol. 56, No. 6


coarse-grained mafic bodies (mafic gneiss and meta-

gabbro) (e.g. Takasu, 1989). The Western Iratsu and

Higashi-akaishi bodies are two of the largest mafic and

ultramafic bodies, respectively (Fig. 1a and b).

Western Iratsu bodyThe Western Iratsu body mainly consists of mafic

gneiss (garnet amphibolite and eclogite) and impuremarble, but surface geological mapping (Kugimiya &

Takasu, 2002; this study) and the core data from a

2300 m borehole (S-7 site) drilled in the years 1967–

1968 by the Metallic Minerals Exploration Agency of

Japan (Supplementary Data Fig. 1; supplementary data

are available for downloading at http://www.petrology.

oxfordjournals.org) have shown that this mafic–calcareous sequence is locally intercalated with ultra-

mafic rocks as a result of pre-DS folding (Fig. 1c). A thick

marble layer within the mafic gneiss contains abundant

mafic fragments and minor metachert layers, suggest-

ing that the provenance of the mafic–calcareous do-

main is the sedimentary facies typically developed on aseamount slope (Kugimiya & Takasu, 2002; Endo,

2010). The ultramafic domain consists of alternating

layers of amphibole-rich metasomatic rocks and a dun-

ite–wehrlite suite. Orthopyroxene-rich rocks highlighted

in this study occur in close association with dunite.

A three-stage metamorphic evolution has been

inferred from the slab-derived mafic–calcareous do-

main (Endo et al., 2009, 2012; Endo, 2010): (1) amphibo-

lite-facies stage (M1) that evolved from 0�9 GPa, 590�Cto 1�2 GPa, 660�C followed by cooling (i.e. a counter-

clockwise P–T path); (2) eclogite-facies stage (M2) with

peak-P conditions of 1�8 GPa, 510–560�C and subse-

quent peak-T conditions of 1�6 GPa, 620�C; (3) epidote–

amphibolite-facies stage (M3: �550�C, 0�8 GPa) that is

the same stage as the regional metamorphic zonation(albite–biotite and oligoclase–biotite zones; Fig. 1a).

The M1 stage is related to an Early Cretaceous

(c. 116 Ma) hot subduction event immediately after the

onset of subduction (Endo et al., 2009, 2012). The

later M2–M3 stages are associated with progressive

warming (increasing thermal gradient) of the subduc-

tion zone and related to a Late Cretaceous (89–85 Ma)event that probably took place just before the Izanagi–

Pacific ridge subduction (Aoya et al., 2003; Wallis et al.,

2009).

W. Iratsu body

1 km

Besshi UnitBesshi UnitBesshi Unit

Eclogite UnitEclogite UnitEclogite Unit

Gongen stream

Higashi-akaishi peridotite body

IR2

IR1

Ultramafic domainMafic domainMarble

MetaserpentiniteMafic, peliticand quartz schists

(c)

10 km

Besshi Unit

Kiyomizu Tectonic Line

Median Tectonic Line(a)

WI

Oboke Unit Recumbent Ds synform

Sanbagawa beltin Shikoku Island

Pre-DsSyn-DsPost-Ds (Du)

Unit boundary

Mafic gneiss Metagabbro Peridotite Serpentinite block

N

HA

Eclogite UnitEclogite UnitEclogite Unit

(b)

1000

2000 m

-1000

Mt. Higashi-akaishi

AABB

S-7

Besshi Unit

Eclogite Unit

Oboke Unit

HAWI

AA

BB

Sea level

Metamorphic grade Oligoclase-biotite zone Albite-biotite zone Garnet zone Chlorite zone

Fig. 1. (a) Tectono-metamorphic map of the Sanbagawa belt in central Shikoku (Aoya et al., 2013a). HA, Higashi-akaishi body; WI,Western Iratsu body. (b) Cross-section along the A–B line (Aoya et al., 2013b). S-7 borehole is also projected perpendicular to theline of section. (c) Simplified geological map of the study area. Locations of samples used in this study (IR1, IR2) are indicated byopen stars.

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FIELD RELATIONSHIPS AND SAMPLINGSTRATEGY

The sample suite used in this study was collected from

a single outcrop (IR1) in the ultramafic domain (Figs 1c

and 2a). This outcrop is composed of coarse-grained

massive dunite, harzburgite and orthopyroxenite, and

corresponds to a low-strain region of the exhumation-related deformation in the Western Iratsu body.

Unfortunately, dunite in this outcrop is almost com-

pletely altered to lizardite serpentinite. The harzburgite

(IR1-h) occurs as a transition zone between dunite and a

>2 m thick layer of orthopyroxenite (IR1-o) (Fig. 2b).

There is no outcrop between the orthopyroxenite and

mafic domains (garnet amphibolite), but antigorite-richshear zones are observed in the orthopyroxenite close

to the contact with the mafic domain. Orthopyroxene–

tremolite–antigorite schist (IR1-s) was collected from

the shear zone. Dunite in the dunite–harzburgite–ortho-

pyroxene sequence is cut by numerous orthopyroxene-

rich veins (IR1-v) (Fig. 2c). These veins can be classified

into two types; one type is represented by thick (�10 cm

wide) symmetrically zoned veins with a tremolite-rich

central zone, and the other type is represented by thin(less than 2 cm wide) orthopyroxene–chlorite veins. The

thin veins branch from the thick veins (Fig. 2d), and are

texturally indistinguishable from the marginal zone of

the thick zoned veins. The branching pattern of the thin

veins linking to a thick vein suggests that fluid flow was

from the mafic domain into the dunite (Fig. 2a and d).As a reference material, dunite that lacked any clear

signs of metasomatism was also collected from a separ-

ate outcrop (IR2) in the Western Iratsu body (Fig. 1c).

The IR2 outcrop is free from orthopyroxene-bearing

rocks and veins. Dunite in this outcrop is moderately

Fig. 2. Field photographs of outcrop IR1 showing the sampling locations. Hammer (40 cm long) or clinometer (7 cm wide) for scaleis shown in each photograph. (a) A panoramic view of the outcrop. (b) Opx1L-rich rocks (harzburgite and orthopyroxenite layers).(c) Opx1V-rich veins in dunite. (d) Close-up view of a thick symmetrically zoned vein and associated thin veins.



serpentinized (i.e. transitional to antigorite schist), butthe degree of lizardite/chrysotile serpentinization is very

weak.

ANALYTICAL METHODS

To display the texture of the orthopyroxene-bearing

rocks, composite element maps of whole thin sections

were acquired by micro X-ray fluorescence (m-XRF)using an M4 Tornado m-XRF spectrometer at the

Geological Survey of Japan (GSJ), AIST. The operating

conditions were 30 kV accelerating voltage, a probe cur-rent of 530mA and 25 mm resolution with an acquisition

time of 5 ms per point. The results are presented in

Fig. 3a–e.

X-ray mapping and quantitative major element ana-

lyses of minerals were carried out using a JEOL

JXA8800 electron microprobe (EMP) at GSJ. X-ray map-ping was conducted with an accelerating voltage of

20 kV, a sample current of 100 nA and an acquisition

time of 20 ms per pixel. The element maps are pre-

sented in Fig. 3f–i. Conditions for quantitative analyses

Fig. 3. m-XRF composite element maps and EMP element maps. (a) CrþNiþCa map of orthopyroxenite (IR1-o). Olivine grains areoutlined in white. It should be noted that anhedral Opx1L grains display euhedral oscillatory zoning, and interstitial spaces are filledby amphiboleþ chlorite. (b) CrþNiþCa map of half of the thick symmetrically zoned vein (IR1-v) shown in Fig. 2d. Tremolite occursonly in the central zone; ‘Dol’ and ‘Serp’ indicate alteration-related dolomite and serpentine veins, respectively. (c, d)CrþFeþCaþS maps showing the distribution of pentlandite in IR1-o and IR1-v. (e) CaþFe map of orthopyroxene–tremolite–antig-orite schist (IR1-s). Black arrows indicate shear bands. (f–h) Al, Ni and Cr maps of the area indicated by the red rectangle in (a).Numbers indicate NiO content (wt %) of olivine. (i) Al map of the area indicated by the red rectangle in (f).



were 15 kV accelerating voltage, 12 nA sample current

and a beam diameter of 2–5 mm. Counting times were

40–50 s on peak and 20–25 s on background for F, Cl

and Ni, and 20–30 s on peak and 10–15 s on background

for other elements. Detection limits for F and Cl are 0�07and 0�01 wt %, respectively. Natural and synthetic sili-

cates and oxides were used as standards. Back-scat-

tered electron (BSE) images were used to evaluate

compositional heterogeneity in each mineral and to

determine analysis locations. The ferric iron content in

Cr-spinel was calculated based on charge balance

assuming perfect stoichiometry. All Fe was assumed tobe Fe2þ for the other minerals. Representative analyses

are reported in Tables 1–4.

In situ trace element compositions of minerals were

determined by laser ablation inductively coupled

plasma mass spectrometry (LA-ICP-MS) at Kanazawa

University, following the procedure described byMorishita et al. (2005). The LA-ICP-MS system consists

of a 193 nm ArF excimer laser (MicroLas laser ablation

system) and an Agilent 7500s quadrupole ICP-MS sys-

tem. Analyses were performed using spot diameters of

110–120mm at a pulse frequency of 5–6 Hz and energy

density of 8 J cm–2. The location of the analysis spots ineach mineral was carefully selected with a petrographic

microscope to avoid both fluid and solid inclusions and

cracks (Supplementary Data Fig. 2). NIST 612 and 614

glass standards were measured every five laser abla-

tion runs on the sample. Acquisition times were 50 s for

background (carrier gas) followed by 60 s for laser abla-

tion on sample. NIST 612 glass (Pearce et al., 1997) wasused for calibration, and 29Si was used as the internal

standard. Data reduction was performed off-line follow-

ing the method proposed by Longerich et al. (1996). The

signal of each analysis was carefully evaluated during

data reduction (Supplementary Data Fig. 2). Trace elem-

ent compositions and associated errors determined for

the NIST 614 glass are shown in Supplementary Data

Table 1. The Cr and Ni contents of orthopyroxene deter-

mined by EMP and LA-ICP-MS analyses are consistent(Supplementary Data Fig. 3). Representative analyses of

minerals are listed in Tables 5 and 6.

Identification of serpentine species was made by

Raman spectroscopy at Nagoya University, using a

Nicolet Almega XR dispersive Raman spectrometer

with a 532 nm Nd-YAG laser. Representative Raman

spectra of antigorite and lizardite from the IR1 outcropare given in Supplementary Data Fig. 4.

Pseudosection modelling was carried out using

Perple_X 6.6.9 (Connolly, 2009) and the thermodynamic

dataset of Holland & Powell (1998, updated in 2002).

Solid-solution models used in this study are the same

as those used by Padron-Navarta et al. (2013).Abbreviations of minerals used in this contribution fol-

low those of Whitney & Evans (2010).

PETROGRAPHY AND MINERAL MAJORELEMENT CHEMISTRY

Dunite (IR2, IR1)Primary anhydrous minerals in dunite in the Western

Iratsu body are olivine and Cr-spinel. Dunite in the IR2

outcrop is transitional to antigorite schist, but least af-

fected by lizardite/chrysotile serpentinization. Olivine

occurs as (1) coarse-grained porphyroclasts and (2)

fine-grained neoblasts in association with platy antigor-ite crystals. This microstructure is very similar to that of

antigorite-bearing dunite in the Higashi-akaishi body

(Mizukami & Wallis, 2005; Wallis et al., 2011). A distinct

feature of the Western Iratsu dunite is the presence of

oriented Cr-spinel lamellae in the porphyroclastic oliv-

ine (Fig. 4a). The two modes of olivine occurrence are

compositionally indistinguishable and are characterizedby high Mg# [¼ Mg/(Mgþ Fe) on an atomic

basis¼ 0�918–0�928] and NiO content (0�34–0�43 wt %),

which overlap with values for olivine in the Higashi-

akaishi dunite (Hattori et al., 2010) (Fig. 5a). The com-

position of primary Cr-spinel is characterized by high

Cr# [¼Cr/(CrþAl)¼ 0�72–0�73] and low TiO2 content(<0�05 wt %), which is also indistinguishable from that

in the Higashi-akaishi dunite (Hattori et al., 2010)

(Fig. 6a–c). Cr-spinel grains are surrounded by magnet-

ite fringes with a narrow intermediate zone with ferrit-

chromite composition (Fig. 6d).

Dunite in the IR1 outcrop was intruded by silicate

veins as described below, and was also subjected tolate-stage alteration to lizardite serpentinite.

Harzburgite (IR1-h) and orthopyroxenite (IR1-o)Orthopyroxene in the transitional harzburgite to ortho-

pyroxenite layers displays subhedral to anhedral stoutcrystals ranging from 2 to 20 mm in length, which are

sparsely distributed in an olivine-rich matrix

Table 1: Electron microprobe analyses (wt %) of olivine

Sample: IR2 IR1h IR1h IR1o IR1o IR1v IR1vreac.f. max Ni thin v. thick v.

SiO2 40�28 39�55 39�69 39�48 39�19 39�74 39�28TiO2 b.d.l. b.d.l. b.d.l. b.d.l. 0�05 b.d.l. b.d.l.Al2O3 b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Cr2O3 b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.FeO 7�75 9�86 10�25 11�06 11�23 9�85 10�66MnO 0�15 0�14 0�13 0�13 0�10 0�10 0�13MgO 51�12 49�91 50�06 49�20 49�48 50�50 48�94CaO b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.NiO 0�38 0�46 0�50 0�56 0�72 0�47 0�60Total 99�68 99�92 100�63 100�43 100�77 100�66 99�61O 4 4 4 4 4 4 4Si 0�98 0�98 0�97 0�97 0�97 0�97 0�98Ti b.d.l. b.d.l. b.d.l. b.d.l. 0�00 b.d.l. b.d.l.Al b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Cr b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Fe 0�16 0�20 0�21 0�23 0�23 0�20 0�22Mn 0�00 0�00 0�00 0�00 0�00 0�00 0�00Mg 1�86 1�83 1�83 1�81 1�82 1�84 1�81Ca b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Ni 0�01 0�01 0�01 0�01 0�01 0�01 0�01Sum 3�01 3�02 3�02 3�02 3�03 3�02 3�02Mg# 0�922 0�900 0�897 0�888 0�887 0�902 0�891

b.d.l., below detection limit.









(harzburgite) (Figs 2b and 4b), or forms granoblastic ag-

gregates (orthopyroxenite) (Figs 3a and 4c). These

orthopyroxene crystals (Opx1L, ‘L’ denotes layers)

contain abundant olivine crystals with a lobate outline

(Figs 3a and 4b, c), euhedral Cr-spinel grains, and

multiphase solid inclusions (MSI) consisting of AmpþChlþPhlþ Ilm (Fig. 4c). Elongate MSI are aligned paral-

lel to the c-axis of host Opx1L (Supplementary Data

Fig. 2). Opx1L grains exhibit oscillatory zoning in Al, Cr

and Ni with euhedral outlines (Fig. 3a, f–h) and are

Table 2: Electron microprobe analyses (wt %) of orthopyroxene

Sample: IR1o IR1o IR1o IR1o IR1h IR1h IR1v IR1v IR1v IR1v IR1sOpx1L Opx1L Opx1L Opx2 Opx1L Opx2 Opx1V Opx1V Opx1V Opx1V Opx2core high Ni rim thin v. cent.z. int.z. mar.z.

SiO2 55�83 55�95 55�35 56�55 56�02 57�03 56�75 56�52 56�21 56�34 56�46TiO2 0�05 0�06 0�09 b.d.l. 0�05 b.d.l. b.d.l. 0�02 0�02 0�02 b.d.l.Al2O3 0�81 0�81 1�15 b.d.l. 0�67 b.d.l. 0�03 0�02 0�02 0�02 b.d.l.Cr2O3 0�57 0�35 0�18 0�02 0�38 0�02 0�04 0�03 0�03 0�03 0�02FeO 6�49 6�83 7�80 7�85 7�06 6�91 6�93 7�65 7�58 7�57 8�76MnO 0�13 0�12 0�16 0�18 0�14 0�17 0�15 0�20 0�19 0�18 0�31MgO 35�05 34�92 34�23 34�75 34�61 35�91 35�91 35�04 35�02 34�99 33�85CaO 0�71 0�78 0�72 0�08 0�60 0�05 0�09 0�08 0�07 0�07 0�04Na2O b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.NiO 0�22 0�34 0�15 0�07 0�15 0�07 0�07 0�09 0�08 0�08 0�04Total 99�87 100�17 99�84 99�50 99�69 100�16 99�97 99�65 99�22 99�30 99�48O 6 6 6 6 6 6 6 6 6 6 6Si 1�94 1�94 1�93 1�97 1�95 1�97 1�97 1�97 1�97 1�97 1�98Ti 0�00 0�00 0�00 b.d.l. 0�00 b.d.l. b.d.l. 0�00 0�00 0�00 b.d.l.Al 0�03 0�03 0�05 b.d.l. 0�03 b.d.l. 0�00 0�00 0�00 0�00 b.d.l.Cr 0�02 0�01 0�01 0�00 0�01 0�00 0�00 0�00 0�00 0�00 0�00Fe 0�19 0�20 0�23 0�23 0�21 0�20 0�20 0�22 0�22 0�22 0�26Mn 0�00 0�00 0�00 0�01 0�00 0�01 0�00 0�01 0�01 0�01 0�01Mg 1�82 1�81 1�78 1�81 1�80 1�85 1�85 1�82 1�83 1�82 1�77Ca 0�03 0�03 0�03 0�00 0�02 0�00 0�00 0�00 0�00 0�00 0�00Ni 0�00 0�00 0�00 0�00 0�00 0�00 0�00 0�00 0�00 0�00 0�00Na b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Sum 4�03 4�03 4�03 4�02 4�02 4�03 4�03 4�02 4�02 4�02 4�02Mg# 0�906 0�901 0�887 0�888 0�898 0�903 0�903 0�891 0�892 0�892 0�874

Table 3: Electron microprobe analyses (wt %) of amphibole

Sample: IR1o IR1o IR1o IR1o IR1o IR1o IR1v IR1sintersti. intersti. intersti. MSI.Opx MSI.Opx MSI.Ol cent.z.

SiO2 49�26 52�85 57�00 49�63 53�34 48�65 56�41 57�33TiO2 0�35 0�19 b.d.l. 0�31 0�10 0�46 0�02 0�02Al2O3 6�10 2�89 0�28 6�35 2�51 6�64 0�46 0�08Cr2O3 0�47 0�48 0�03 1�14 0�78 0�87 0�05 b.d.l.FeO 3�77 2�76 2�12 2�93 2�34 3�50 2�01 2�28MnO 0�05 0�07 0�08 0�05 0�02 0�05 0�06 0�15MgO 21�24 22�83 23�75 21�14 22�93 21�07 23�68 24�16CaO 11�84 12�25 12�40 12�02 12�26 12�23 12�44 12�59Na2O 2�32 1�56 0�57 2�25 1�32 2�49 0�76 0�24K2O 0�90 0�49 0�10 0�48 0�29 0�56 0�12 0�04NiO 0�13 0�12 0�11 0�32 0�14 0�17 0�12 n.a.F b.d.l. 0�17 n.a. n.a. n.a. n.a. b.d.l. n.a.Cl b.d.l. b.d.l. n.a. n.a. n.a. n.a. b.d.l. n.a.Total 96�43 96�37 96�44 96�62 96�03 96�69 96�13 96�89O 23 23 23 23 23 23 23 23Si 7�04 7�45 7�91 7�05 7�52 6�94 7�86 7�91Ti 0�04 0�02 b.d.l. 0�03 0�01 0�05 0�00 0�00Al 1�03 0�48 0�05 1�06 0�42 1�12 0�08 0�01Cr 0�05 0�05 0�00 0�13 0�09 0�10 0�01 b.d.l.Fe 0�45 0�33 0�25 0�35 0�28 0�42 0�23 0�26Mn 0�01 0�01 0�01 0�01 0�00 0�01 0�01 0�02Mg 4�53 4�80 4�91 4�48 4�82 4�48 4�92 4�97Ca 1�81 1�85 1�84 1�83 1�85 1�87 1�86 1�86Na 0�65 0�43 0�16 0�62 0�36 0�69 0�21 0�07K 0�17 0�09 0�02 0�09 0�05 0�10 0�02 0�01Ni 0�02 0�01 0�01 0�04 0�02 0�02 0�01 n.a.Sum 15�78 15�52 15�15 15�67 15�42 15�79 15�20 15�11Mg# 0�910 0�937 0�952 0�928 0�946 0�915 0�955 0�950

n.a., not analysed.





Table 5: LA-ICP-MS analyses (ppm, average 6 1r) of orthopyroxene and olivine

Sample: IR1o IR1o IR1v IR1v IR1v IR2 IR1oOpx1L Opx2 Opx1V Opx1V Opx1V Ol Ol

cent.z. int.z. mar.z.

Li 3�36 (0�36) 3�47 (1�88) 0�88 (0�19) 0�80 (0�25) 1�01 (0�31) 2�55 (0�19) 3�13 (0�17)B 0�94 (0�11) 3�58 (0�54) 1�27 (0�48) 1�12 (0�22) 1�25 (0�23) 2�17 (0�72) 1�63 (0�13)Sc 5�36 (0�59) 1�96 (0�20) 1�42 (0�15) 1�25 (0�88) 1�26 (0�05) 2�55 (0�13) 1�78 (0�04)Ti 294 (29) 100 (10) 117 (14) 121 (10) 121 (4) 3�99 (0�38) 28�9 (4�2)V 15�6 (2�3) 1�74 (0�67) 1�12 (0�55) 0�88 (0�21) 0�96 (0�05) 0�11 (0�04) 1�98 (0�53)Cr 3247 (849) 82�2 (13�9) 171 (37) 186 (62) 207 (24) 6�68 (2�98) 352 (132)Co 98�0 (19�3) 66�3 (2�5) 73�3 (5�5) 69�3 (1�4) 69�1 (1�5) 149 (2) 213 (6)Ni 2072 (516) 589 (55) 789 (317) 665 (87) 654 (30) 3090 (44) 4830 (80)Rb <0�010 0�09 (0�07) <0�015 <0�015 0�039 <0�010 <0�010Sr 0�24 (0�04) 0�22 (0�14) 0�11 (0�06) 0�040 (0�006) 0�040 (0�005) 0�002 (0�000) 0�020 (0�014)Y 0�24 (0�06) 0�070 (0�006) 0�089 (0�004) 0�096 (0�008) 0�098 (0�009) 0�003 0�005 (0�000)Zr 0�22 (0�06) 0�010 (0�006) 0�008 0�010 (0�006) 0�010 (0�006) 0�007 (0�002) 0�012 (0�002)Nb 0�024 (0�006) 0�003 (0�002) 0�005 (0�001) 0�008 (0�003) 0�007 (0�004) 0�008 (0�001) 0�018 (0�006)Cs <0�006 1�14 (1�02) 0�15 0�014 (0�012) 0�15 (0�15) <0�005 <0�005Ba 0�034 (0�030) 0�044 (0�011) 0�036 0�006 (0�004) 0�014 (0�006) 0�029 (0�006) 0�027 (0�020)La 0�007 (0�003) 0�005 (0�001) 0�010 (0�002) 0�016 (0�003) 0�010 (0�002) 0�001 (0�000) 0�005 (0�001)Ce 0�036 (0�009) 0�016 (0�003) 0�034 (0�010) 0�042 (0�005) 0�030 (0�007) 0�002 (0�001) 0�012 (0�003)Pr 0�006 (0�002) 0�002 (0�000) 0�004 (0�001) 0�005 (0�001) 0�003 (0�001) <0�001 0�001 (0�000)Nd 0�045 (0�012) 0�010 (0�001) 0�020 (0�003) 0�025 (0�002) 0�017 (0�005) <0�004 <0�004Sm 0�025 (0�007) 0�003 <0�012 0�008 (0�000) 0�007 (0�001) <0�004 <0�004Eu 0�008 (0�002) 0�002 (0�000) 0�002 (0�000) 0�002 (0�000) 0�002 (0�000) <0�001 <0�001Gd 0�033 (0�012) 0�006 0�012 (0�003) 0�010 (0�003) 0�009 (0�001) <0�006 <0�006Tb 0�006 (0�002) <0�001 <0�002 0�002 0�002 (0�000) <0�001 <0�001Dy 0�042 (0�012) 0�011 (0�001) 0�014 (0�000) 0�016 (0�002) 0�016 (0�002) <0�003 <0�004Ho 0�009 (0�002) 0�003 (0�000) 0�003 (0�000) 0�004 (0�000) 0�004 (0�000) <0�001 <0�001Er 0�027 (0�007) 0�009 (0�001) 0�011 (0�001) 0�012 (0�003) 0�013 (0�001) <0�003 <0�003Tm 0�005 (0�001) 0�002 0�002 (0�001) 0�003 (0�000) 0�003 (0�000) <0�001 <0�001Yb 0�035 (0�010) 0�017 (0�003) 0�022 (0�002) 0�021 (0�003) 0�024 (0�003) <0�006 <0�005Lu 0�006 (0�002) 0�003 (0�000) 0�004 (0�001) 0�004 (0�000) 0�004 (0�000) <0�001 <0�001Hf 0�010 (0�003) <0�003 <0�010 <0�010 <0�010 <0�006 <0�007Ta <0�002 <0�001 <0�002 <0�002 <0�003 <0�002 <0�002Pb 0�15 (0�10) 0�115 (0�058) 0�035 (0�026) 0�038 (0�009) 0�041 (0�018) 0�018 (0�008) 0�041 (0�022)Th 0�006 (0�003) 0�013 0�011 (0�001) 0�050 (0�016) 0�048 (0�028) <0�002 0�006 (0�002)U 0�003 (0�001) 0�004 (0�002) 0�003 0�009 (0�002) 0�008 (0�002) <0�003 0�004 (0�001)

Table 4: Electron microprobe analyses (wt %) of chlorite, phlogopite and antigorite

Sample: IR1o IR1o IR1o IR1v IR1v IR1o IR2 IR1o IR1sMineral: Chl Chl Chl Chl Chl Phl Atg Atg Atg

intersti. MSI.Opx MSI.Ol thin v. thick v. MSI.Opx

SiO2 32�69 32�93 31�72 31�78 31�80 41�14 43�13 40�55 42�15TiO2 b.d.l. b.d.l. b.d.l. b.d.l. 0�02 0�27 b.d.l. 0�02 b.d.l.Al2O3 13�34 14�34 13�05 12�34 12�63 11�72 0�35 3�18 1�47Cr2O3 1�98 1�48 1�95 3�33 2�65 0�26 0�07 0�57 0�16FeO 4�08 3�44 4�30 3�64 3�83 2�85 2�51 4�25 4�44MnO b.d.l. b.d.l. b.d.l. 0�02 0�01 0�02 0�05 0�02 0�05MgO 33�04 32�75 33�55 33�80 33�38 26�96 39�24 36�89 37�45CaO b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Na2O n.a. n.a. n.a. n.a. n.a. 0�51 n.a. n.a. n.a.K2O n.a. n.a. n.a. n.a. n.a. 9�53 n.a. n.a. n.a.Total 85�13 84�94 84�57 84�91 84�32 93�26 85�35 85�48 85�72O 14 14 14 14 14 11 116 116 116Si 3�18 3�18 3�12 3�11 3�13 2�96 33�89 32�21 33�32Ti b.d.l. b.d.l. b.d.l. b.d.l. 0�00 0�01 b.d.l. 0�01 b.d.l.Al 1�53 1�63 1�51 1�43 1�47 0�99 0�33 2�98 1�38Cr 0�15 0�11 0�15 0�26 0�21 0�02 0�05 0�36 0�10Fe 0�33 0�28 0�35 0�30 0�32 0�17 1�65 2�82 2�94Mn b.d.l. b.d.l. b.d.l. 0�00 0�00 0�00 0�03 0�02 0�04Mg 4�79 4�72 4�91 4�94 4�90 2�89 45�97 43�69 44�13Ca b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l.Na n.a. n.a. n.a. n.a. n.a. 0�07 n.a. n.a. n.a.K n.a. n.a. n.a. n.a. n.a. 0�87 n.a. n.a. n.a.Sum 9�98 9�92 10�04 10�04 10�00 7�97 81�92 82�09 81�91Mg# 0�935 0�944 0�933 0�943 0�939 0�944 0�965 0�939 0�938



locally crosscut by Al-poor orthopyroxene (Opx2:

Al2O3< 0�04 wt %, Al< 0�002 a.p.f.u./6O)þAtg 6 Tr 6

Chl 6 Ilm (Fig. 3i). Opx1L in IR1-o shows the following

compositional range: Mg# 0�881–0�918, Al2O3 0�45–1�2 wt %; Al¼0�02–0�05 a.p.f.u., Cr2O3 0�06–0�65 wt %,

CaO 0�26–1�0 wt % and NiO 0�08–0�40 wt % (Fig. 7). The

zoning profile in Opx1L is characterized by an overall

trend of decreasing Mg# and Cr2O3 content from the

core to the rim (Fig. 7c). The Ni zoning in Opx1L shows

two Ni-rich zones: a homogeneous core and a narrowouter zone, which shows the highest Ni contents

(Fig. 7c). Relatively low Ni halos (with comparable

Ni content to the Ni-poor zone) are observed around

olivine inclusions in Opx1L (Fig. 3g). Small (�50 mm)

pentlandite grains are distributed throughout the Ni-

poor rim of Opx1L (Fig. 3c and g).

Olivine in Opx1L-rich rocks contains numerous ori-ented Cr-spinel lamellae (�20mm� 2 mm) and unor-

iented AmpþChl composite inclusions. Cr-spinel

lamellae in olivine in Opx1L-rich rocks are more abun-

dant than those in dunite (Fig. 4d). Exceptionally large

spinel lamellae sufficient for EMP analysis have Cr-

magnetite compositions (Fig. 6b), but the exact com-position of smaller brown transparent lamellae has not

been determined. Olivine grains enclosed in Opx1L are

mostly monocrystalline and commonly have the same

extinction orientation for neighboring inclusions under

crossed polars (Fig. 4b). Rare polycrystalline olivine

enclosed in Opx1L crystals is associated with

AmpþChl aggregates along the grain boundaries ofolivine (Fig. 3f). Relict olivine grains in Opx1L-rich rocks

from the IR1 outcrop show compositional trends with

increasing NiO content (up to 0�74 wt %) and decreasing

Mg# (0�880–0�895 for orthopyroxenite) with increasing

amount of Opx1L (Fig. 5a). Olivine associated with pent-

landite and the Ni-poor rim of Opx1L have lower NiOcontents (down to 0�43 wt %) (Figs 3g and 5a).

Euhedral Cr-spinel grains included in the core of

Opx1L (Fig. 6e) are chromite showing a weak zoning

with increasing TiO2 content towards the rim (Fig. 6g).

Cr-spinel grains enclosed in the outer zones of Opx1L

are commonly surrounded by chlorite and variably oxi-

dized to ferritchromite/Cr-magnetite in their rims (Fig. 6fand h).

Amphibole occurs in three modes: (1) filling intersti-

tial spaces between orthopyroxene crystals (Fig. 3a and

i); (2) constituting MSI in Opx1L; (3) forming amphi-

bole–chlorite aggregates that are included in monocrys-

talline olivine or fill grain boundaries of polycrystallineolivine (Fig. 3f). These amphibole crystals show com-

positional zoning from hornblende (edenite to magne-

siohornblende according to Leake et al., 1997) in the

Table 6: LA-ICP-MS analyses (ppm, average 6 1r) of amphibole, phlogopite and chlorite

Sample: IR1o IR1o IR1v IR1o IR1o IR1vHbl Tr Tr Phl Chl Chlintersti. intersti. cent.z. MSI.Opx intersti. int.z.

Li 2�28 (0�20) 1�79 1�43 (0�12) 26�3 0�55 (0�35) 0�29 (0�17)B 13�5 (1�3) 8�27 8�43 (1�61) 2�42 1�50 (0�03) 3�99 (0�03)Sc 47�4 (1�9) 26�3 10�8 (1�2) 4�66 9�50 (0�05) 5�69 (0�14)Ti 601 (31) 298 113 (25) 773 128 (4) 97�5 (7�7)V 74�3 (10�2) 25�7 5�96 (0�05) 51�7 79�9 (1�1) 65�5 (0�8)Cr 3344 (1351) 892 257 (109) 10229 7815 (424) 13867 (17)Co 43�1 (5�2) 33�4 29�7 (1�4) 71�4 61�9 (1�3) 57�9 (2�0)Ni 1282 (157) 973 899 (48) 3331 1833 (7) 2218 (130)Rb 3�69 (0�31) 0�39 0�12 (0�03) 227 4�14 (1�47) 0�047 (0�002)Sr 278 (19) 185 308 (48) 4�05 0�14 (0�04) 0�14 (0�09)Y 11�6 (0�5) 5�39 3�93 (0�47) 0�079 0�003 (0�001) <0�002Zr 33�0 (7�1) 3�4 0�35 (0�12) 0�23 0�061 (0�009) 0�22 (0�03)Nb 1�18 (0�21) 0�14 0�041 (0�040) 0�22 0�079 (0�016) 0�21 (0�00)Cs 0�38 (0�33) 0�06 0�20 (0�12) 5�5 0�39 (0�34) 0�14 (0�08)Ba 44�5 (3�7) 4�48 0�19 (0�03) 1018 13�2 (5�9) 0�19 (0�09)La 7�30 (1�04) 1�2 0�36 (0�08) 0�006 <0�001 <0�002Ce 30�0 (3�1) 5�62 2�52 (0�56) 0�024 0�004 (0�002) 0�001 (0�000)Pr 4�16 (0�33) 0�93 0�54 (0�11) 0�003 <0�001 <0�001Nd 19�8 (1�3) 4�95 3�40 (0�68) 0�01 <0�003 <0�005Sm 4�71 (0�20) 1�47 1�25 (0�21) <0�009 <0�004 <0�010Eu 1�28 (0�06) 0�43 0�28 (0�04) 0�01 <0�002 <0�003Gd 3�77 (0�17) 1�35 1�17 (0�17) 0�06 <0�006 <0�008Tb 0�46 (0�02) 0�19 0�15 (0�02) <0�004 <0�002 <0�003Dy 2�63 (0�10) 1�14 0�90 (0�13) <0�006 <0�005 <0�007Ho 0�45 (0�02) 0�2 0�15 (0�02) <0�002 <0�002 <0�002Er 1�12 (0�05) 0�53 0�38 (0�05) <0�004 <0�003 <0�007Tm 0�14 (0�01) 0�072 0�050 (0�007) <0�002 <0�002 <0�003Yb 0�86 (0�04) 0�44 0�32 (0�05) <0�010 <0�004 <0�006Lu 0�10 (0�00) 0�054 0�039 (0�006) <0�002 <0�002 <0�002Hf 1�99 (0�45) 0�14 0�064 (0�023) 0�038 <0�007 <0�005Ta 0�12 (0�04) 0�009 0�007 0�045 <0�002 <0�003Pb 7�57 (2�56) 2�36 1�35 (0�29) 0�38 0�11 (0�02) 0�018Th 0�15 (0�05) 0�032 0�011 (0�008) 2�96 0�58 (0�20) <0�003U 0�025 (0�020) 0�004 0�005 (0�000) 1�44 0�046 (0�017) 0�010 (0�001)



core to tremolite in the rim (Figs 3i and 8a). The horn-

blende contains up to 0�5 wt % TiO2 (Fig. 8b) and 1�9 wt% Cr2O3 (Fig. 8c). Chlorite in MSI in Opx1L is richer in Al

compared with this mineral occurring in interstitial

spaces or as discrete inclusions in Opx1L (Fig. 9a).

Antigorite contains appreciable amounts of Tschermakcomponent (Al¼ 1�9–3�6 a.p.f.u., Cr¼0�10–0�85 a.p.f.u.,

Mg#¼ 0�932–0�944) (Fig. 9b and c).

Fig. 4. Photomicrographs of thin sections. (a) Olivine with Cr-spinel lamellae in dunite (IR2). Plane-polarized light (PPL). (b) Opx1Lcrystal in harzburgite (IR1-h) showing a highly irregular interface with olivine. Olivine in the matrix is altered to lizardite (Lz) andmagnetite. Cross-polarized light (CPL). (c) Anhedral Opx1L crystals in orthopyroxenite (IR1-o). Opx1L crystals have a poikilitic tex-ture enclosing irregular-shaped olivine (Ol), euhedral Cr-spinel (Spl), and multi-phase solid inclusions (MSI). CPL. (d) Oriented Cr-magnetite lamellae in olivine in IR1-o. PPL. (e) Elongated Opx1V crystals in the marginal zone of the thick vein (IR1-v). The (001)cleavage of chlorite flakes is parallel to the (100) plane of Opx1V. PPL. (f) Opx2–tremolite–antigorite schist (IR1-s). CPL.



Orthopyroxene-bearing veins (IR1-v)The thick symmetrically zoned vein in dunite (Fig. 2c

and d) is composed of the following three zones: a cen-

tral zone (OpxþTrþChl), an intermediate zone (OpxþChl), and a marginal zone (OpxþOlþChl) (Fig. 3b). Inthin veins and in the marginal zones of thick veins,

elongate euhedral crystals up to 1 cm in length of

Opx1V (‘V’ denotes veins) occur scattered in the olivine-

dominated matrix. The crystal habit of Opx1V is pris-

matic in the [001] direction and tabular on (100).

Compared with Opx1L, Opx1V is significantly lower inAl2O3 (<0�05 wt %; up to 0�002 a.p.f.u. Al), Cr2O3

(<0�07 wt %) and CaO (<0�12 wt %). The NiO content in

Opx1V is relatively high (up to 0�15 wt %) and Mg# de-

creases from 0�913 in the thin vein to 0�888 in the thick

vein (Fig. 7a and b).

Chlorite is intimately associated with Opx1V where

the chlorite (001) plane is parallel to the (100) face ofOpx1V crystals (Fig. 4e). Chlorite grains are partially

replaced by antigorite. Chlorite is rich in Cr and this is

particularly marked in thin veins (Fig. 9a).

The elongated Opx1V crystals contain inclusions of

chlorite, olivine and Cr-spinel, whereas stout Opx1V

crystals in the inner zones of thick veins lack olivine in-

clusions. Pentlandite occurs as disseminated grains

(�50 mm) in the inner (i.e. olivine-absent) zones of theveins (Fig. 3d). Olivine in the marginal zone of the thick

vein is characterized by higher NiO content (up to

0�68 wt %) and lower Mg# (0�885–0�895) compared with

that in thin veins (Fig. 5b).

Tremolite in the central zone occurs as an aggregate

of small (�0�5 mm in length) prismatic crystals and is

compositionally homogeneous (Fig. 8a).

Orthopyroxene–tremolite–antigorite schist(IR1-s)Ductile deformation of the orthopyroxenite during theantigorite-stable stage is localized in shear zones where

Opx1L, olivine and Cr-spinel are completely decom-

posed into coarse-grained Opx2, tremolite, antigorite

and fine-grained magnetite (Figs 3e and 4f).

Constituent minerals (Opx, Atg, Tr and Mag) in IR1-s

are chemically homogeneous. Large orthopyroxenecrystals (former Opx1L) are completely re-equilibrated

to Al-poor orthopyroxene (Opx2: <0�04 wt % Al2O3).

Opx2 in IR1-s is characterized by low Mg# (0�865–0�879)

(Fig. 7a and b) owing to high modal abundance of antig-

orite (Fig. 3e). Acicular tremolite crystals invade ortho-

pyroxene crystal margins (Fig. 4f) and define a

schistosity and shear bands together with platy antigor-ite crystals and stringers of magnetite grains (Fig. 3e).

Tremolite has Mg# ¼ 0�948–0�955 and is close to the

end-member composition (Fig. 8a). Antigorite has a

moderate Tschermak component (Al¼ 0�53–1�89

a.p.f.u., Cr¼ 0�03–0�20 a.p.f.u., Mg# ¼ 0�933–0�943)

(Fig. 9b and c).

MINERAL TRACE ELEMENT CHEMISTRY

OrthopyroxenePrimitive mantle (PM) normalized trace element abun-

dances in Opx1L and Opx2 are shown in Fig. 10a.

Compared with Opx1L, Opx2 has significantly lower

concentrations of Sc, Y, rare earth elements [REE; par-ticularly the middle REE (MREE)], high field strength

elements (HFSE; Nb, Zr, Hf and Ti) and transition metals

(V, Cr and Ni). Chondrite-normalized REE abundances

in Opx1L show a nearly flat pattern from heavy REE

(HREE) to MREE (SmN/YbN¼ 0�67–0�91) and a steep de-

crease from MREE to light REE (LREE) (LaN/SmN¼ 0�11–

0�30) with a slight negative Eu anomaly [EuN/Eu*¼ 0�16–0�23, Eu*¼ (SmNþGdN)/2], whereas those

in Opx2 show a monotonous decrease from HREE to

LREE (Fig. 10b).

In the thick zoned vein (IR1-v), the trace element

compositions of Opx1V are similar in all the three tex-

tural zones (Fig. 10c and d). PM-normalized abundancesof most trace elements (Nb, Sr, Zr, Hf, MREE, Ti, Y,

HREE, Sc, V, Cr and Ni) in Opx1V are significantly lower

OlivineN

iO (w

t %)

Mg#

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0.87 0.88 0.89 0.90 0.91 0.92 0.93 0.94

Olivine mantle array

Increasing Opx1L

Pn

form

atio

n

Higashi-akaishi dunite

Dunite (IR2)

IR1-h (Reac. front)IR1-hIR1-o

Olivine mantle array

NiO

(wt %

)

Mg#

0.2

0.3

0.4

0.5

0.6

0.7

0.87 0.88 0.89 0.90 0.91 0.92 0.93 0.94

IR1-v (thin vein)

IR1-v (thick vein)

Increasing Opx1V

(a)

(b)Olivine

1σ

1σ

Fig. 5. Olivine compositions plotted on Mg# vs NiO diagrams.(a) Olivine in dunite (IR2) and Opx1L-bearing rocks (IR1-h andIR1-o). The olivine mantle array (Takahashi et al., 1987) andcompositional range of olivine in the Higashi-akaishi dunite(Hattori et al., 2010) are also shown. (b) Olivine in Opx1V-richveins (IR1-v).



than in Opx1L but comparable with Opx2 in IR1-o

(Fig. 10a). Chondrite-normalized REE abundances of

Opx1V show a slight enrichment in LREE (LaN/SmN¼0�74–1�4) with a negative Eu anomaly (EuN/Eu*¼0�14–0�39).

OlivineMost incompatible trace elements (other than Li and B)

in olivine are close to or below the detection limits of

LA-ICP-MS analysis. Analysis spots of olivine in ortho-

pyroxenite (IR1-o) were carefully selected to avoid

AmpþChl inclusions as far as possible, but it was im-

possible to avoid Cr-spinel lamellae owing to their small

size and high abundances. Data for Ba, Pb, Th, U andLREE were not considered because the very low, but

nevertheless detectable, concentrations of these elem-

ents in some analyzed spots suggest a contribution

from fluid or Amp inclusions. Because the oriented Cr-

spinel inclusions are thought to be exsolution lamellae

from former Cr-bearing olivine and the Cr-spinel lamel-lae are sufficiently small relative to the spot size of

LA-ICP-MS analyses, measured compositions of olivine

(þ Cr-spinel lamellae) in IR1-o approximate the pre-

existing Cr-bearing olivine composition. Compared with

clear olivine in dunite (IR2), olivine with Cr-spinel lamel-

lae in IR1-o has significantly higher abundances in Ti, Vand Cr (Fig. 11).

AmphiboleHornblende in orthopyroxenite (IR1-o) is an important

contributor to the trace element budget in the rock. PM-

normalized trace element patterns of hornblende show

negative anomalies in Th, U and HFSE (Nb, Zr, Hf and

Ti). Tremolite in IR1-o has much lower trace element

abundances (except for Sr) than hornblende, although

the PM-normalized trace element patterns of tremoliteare similar to those of hornblende (Fig. 12a). Chondrite-

normalized REE abundances in hornblende show a

smooth increase from HREE to LREE with a downward

inflection in La (Fig. 12b).

Tremolite in the thick Opx1V-rich vein (IR1-v) has

trace element abundances of �0�1�PM for Ba, Th, Nb,Zr, Hf and Ti, and �10�PM for Pb and Sr (Fig. 12a). The

low Ba content is a remarkable feature. The REE

Cr

Al

Fe3+

(a)

0

0.1

0.2

0.3

0 60 1200

0.4

0.8

1.2

0 60 120

0

0.4

0.8

1.2

0 0.2 0.4 0.6 0.8 1

Cr/(Cr + Al + Fe3+)

TiO

2 (w

t%)

Higashi-akaishi dunite

Dunite (IR2)

IR1-hIR1-o

(b)

TiO2

YFe3+

Mg#

YFe3+

TiO2

Mg#

A

B

C D

A B C D

Opx1L(core)

IR2

IR1-o IR1-o

Spl (unaltered)

Mag

Ftc Cr-Mag

Cr-Mgt

Mag

μm μm

TiO2 (w

t%)

Mg#

, YFe

3+

(d)

(e) (f)

(g) (h)

Hbl+Chl

600oC

550o C

0

0.1

0.2

0.3

0.4

Mg#

Lamellae in Ol

Ftc

Mag

Cr-Mgt

Ftc

Spl

Spl

Ftc

Core

Core

Core

Rim

Rim

Rim

(c)

Fig. 6. (a) Trivalent cation ratios of zoned Cr-spinel grains in dunite (IR2) and Opx1L-bearing rocks (IR1-h and IR1-o). Compositionalrange of primary Cr-spinel in the Higashi-akaishi dunite (Hattori et al., 2010) is shown for comparison. Dashed lines show the limitsof spinel solid solution in equilibrium with olivine (Mg#¼0�9) at 550�C and 600�C (Sack & Ghiorso, 1991). (b, c) Variation of TiO2

content and Mg# in zoned Cr-spinel vs Cr/(CrþAlþFe3þ). (d) BSE image of zoned Cr-spinel in IR2. (e) BSE image of Cr-spinel grainsincluded in the core of Opx1L shown in Fig. 3h. (f) BSE image of a slightly oxidized Cr-spinel grain in IR1-o. (g, h) Mg#, YFe3þ

[¼Fe3þ/(CrþAlþFe3þ)] and TiO2 profiles across the traverses A–B and C–D in (e) and (f).



patterns of tremolite in IR1-v are characterized by a

steep increase from LREE to MREE (LaN/SmN¼ 0�16–

0�20) followed by a decrease from MREE to HREE

(SmN/YbN¼ 4�2–4�8) with a negative Eu anomaly (EuN/

Eu*¼ 0�16–0�18) (Fig. 12b).

Other mineralsFigure 13 shows PM-normalized trace element abun-

dances of chlorite in IR1-v and some incompatible elem-

ent-rich (i.e. probably former fluid-rich) domains in

IR1-o: an MSI-rich domain within Opx1, an amphibole

(Hbl/Tr)–chlorite aggregate filling the grain boundary of

polycrystalline olivine enclosed in Opx1 (Fig. 3f),phlogopite in an exceptionally large MSI (TrþChlþPhlþ Ilm) in Opx1 (Fig. 3f) and chlorite in a former pore

space (Fig. 3i). Phlogopite is a major host for Rb, Ba, Th

and U. Chlorite is not a significant host for incompatible

trace elements.

PRESSURE–TEMPERATURE ESTIMATES

Formation conditions of Opx1LThe evolution of mineral parageneses in the IR1 outcrop

is summarized in Fig. 14a. Orthopyroxene is stable on

the high-temperature side of the following reactions:

Tlcþ Atg ¼ Opxþ H2O

TlcþOl ¼ Opxþ H2O:

0

0.2

0.4

0.6

0.8

1.0

1.2

1.4

0

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0

0.2

0.86 0.88 0.90 0.92 0.940

0.1

0.86 0.88 0.90 0.92 0.94

Mg# Mg#

(b) (a) Al2O3 (wt%) Cr2O3 (wt%)

IR1-o (Opx1L)IR1-o (Opx2)

IR1-h (Opx1L)IR1-h (Opx2)

IR1-s (Opx2)

IR1-v (thick vein)IR1-v (thin vein)

core rimhigh-Ni

0.88

0.89

0.90

0.91

0.92

0.2

0.4

0.6

0.8

1.0

1.2

0 0.4 0.8 1.2 1.6 2.0 2.4

NiO

Cr2O3

Al2O3

Mg# Mg#

wt%

BA Distance (mm)

(c)

700

800

900

T (o C

)

Orthopyroxene

0

Fig. 7. Compositions of orthopyroxene (Opx1L, Opx1V and Opx2). (a) Mg# vs Al2O3. (b) Mg# vs Cr2O3. (c) Mg#, Al2O3, Cr2O3 andNiO profiles along the A–B profile in Fig. 3g. Apparent temperature is calculated using the Cr–Al in Opx geothermometer (Witt-Eickschen & Seck, 1991).



In addition, the following reaction constrains the high-

temperature limit for the assemblage Opx1Lþchlorite:

Chl ¼ OpxþOlþ Splþ H2O:

The approximate range of Opx1L formation conditions

is constrained to be �650–850�C in the spinel peridotitefacies (Bose & Ganguly, 1995; Fumagalli & Poli, 2005;

Grove et al., 2006). Under antigorite-unstable and chlor-

iteþorthopyroxene-stable conditions, the modal abun-

dance of chlorite solid solution increases with

decreasing temperature, forming low-Al orthopyroxene

at low temperatures. Thus, the relatively high Al content

of Opx1L suggests that its formation temperature isclose to the upper stability limit of chlorite (�800–

850�C) (Fig. 14b).

Opx1L coexists with Cr-spinel and its compos-

itional range is entirely in the applicable range of the

empirical Cr–Al in orthopyroxene geothermometer

(Witt-Eickschen & Seck, 1991). This geothermometer

indicates a general tendency of decreasing temperature

from the core (�880�C) to the rim (�750�C) of each

Opx1L grain (Fig. 7c). Because Opx1L displays an

oscillatory zoning in Cr and Al, the distribution ofthese elements in Opx1 may deviate from chemical

equilibrium to some degree (e.g. Yardley et al., 1991;

Shore & Fowler, 1996). Nevertheless, the calculated

temperature is consistent with the aforementioned

conditions.

[A] N

a +

K (a

pfu/

23O

)

Si (apfu/23O)

(a)

0

0.2

0.4

TiO

2 (w

t%)

0

0.5

1.0

1.5

Cr 2

O3

(wt%

)

0

0.5

6.57.07.58.0

Tr Hbl

Ed

Si (apfu/23O)6.57.07.58.0

(b)

IR1-oIR1-vIR1-s

Tr Hbl

To PrgTo Ed

InterstitialMSI in OlMSI in Opx1L

IR1-o

(c)

Fig. 8. (a) Compositions of amphibole in orthopyroxenite (IR1-o), thick Opx1V-rich vein (IR1-v) and Opx2–tremolite–antigoriteschist (IR1-s) plotted on an Si vs [A]NaþK diagram. (b, c)Variation in TiO2 and Cr2O3 contents of amphibole in IR1-o plot-ted against Si content.

0

1

2

3

4

31 32 33 34

Al+

Cr (

apfu

/116

O)

Si (apfu/116O)

(c) 5 Tschermak exchange (m = 17)

0

0.1

0.2

0.3

1.0 1.2 1.4 1.6 1.8 2.0

IR1-oIR1-s

IR2

0

0.2

0.4

0.6

0.8

1.0

0 1 2 3 4

Cr (

apfu

/116

O)

Al (apfu/116O)

(b)

Cr (

apfu

/14O

)

Al (apfu/14O)

(a)

InterstitialMSI in OlMSI in Opx1L

IR1-o

IR1-v

Thin veinThick vein

Chlorite

Antigorite

Fig. 9. (a) Chlorite compositions plotted on an Al vs Cr diagram.(b, c) Antigorite compositions plotted on Al vs Cr and Si vsAlþCr diagrams.



The estimated temperature range is significantly

higher than the peak conditions of the amphibolite-

facies metamorphism (M1) in the slab-derived mafic do-

main of the Western Iratsu body (Endo et al., 2009,2012) (Fig. 14b).

Formation conditions of Opx1VThe low Al (less than 0�002 a.p.f.u.) and Cr contents of

Opx1V in the presence of chlorite and Cr-spinel suggest

a significantly lower formation temperature of Opx1V

compared with Opx1L. Thus, Opx1V in the ultramafic

domain may have formed coevally with the peak of theamphibolite-facies metamorphism (M1: 660�C and

1�2 GPa) in the mafic domain (Endo et al., 2009, 2012)

(Fig. 14b).

The symmetrical arrangement of mineralogical

zones in the thick Opx1V-rich vein (IR1-v) (Fig. 3b and d)

suggests influx of a Si-rich fluid into a brittle fracture indunite, and the growth of the metasomatic reaction

vein is controlled by diffusion of components from the

fluid to the reaction front (e.g. Bucher, 1998; Markl et al.,

2003). This process is driven by chemical potential

gradients (Korzhinskii, 1959). To model the stability of

mineral assemblages in the thick zoned vein, mSiO2–

mCaO–mAl2O3 pseudosections in the system CaO–FeO–MgO–Al2O3–SiO2–H2O were calculated at 660�C and

1�2 GPa (Fig. 15a and b). Under these P–T conditions,

low-Al orthopyroxene is stable in the OpxþChl field

(Fig. 15b). The observed mineral zones in the

(a)

0.01

0.1

1

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

0.001

0.0001

0.01

0.1

1

10

BaPb

ThU

NbLa

CePr

SrNd

ZrHf

SmEu

TiGd

TbDy

HoY

ErTm

YbLu

ScV

CrCo

Ni

0.001

0.01

0.1

1

10

BaPb

ThU

NbLa

CePr

SrNd

ZrHf

SmEu

TiGd

TbDy

HoY

ErTm

YbLu

ScV

CrCo

Ni0.01

0.1

1


(b)

(c) (d)

IR1-o IR1-o

IR1-v IR1-v

Opx1LOpx2

Marginal zoneIntermediate zoneCentral zone

Opx1L

Opx/C1 chondriteOpx/Primitive mantle

Opx/C1 chondriteOpx/Primitive mantle

Opx1L

Opx1V in thick vein

0.0001

Fig. 10. (a) Primitive mantle (McDonough & Sun, 1995) normalized trace element patterns of Opx1L and Opx2 in orthopyroxenite(IR1-o). (b) C1 chondrite (McDonough & Sun, 1995) normalized rare earth element patterns of Opx1L and Opx2 in IR1-o. (c) Traceelement patterns of Opx1V in thick vein (IR1-v). (d) REE patterns of Opx1V in IR1-v.

Li B Nb Zr Ti Y Sc V Cr Co Ni

0.001

0.0001

0.01

0.1

1

10

Sam

ple/

Prim

itive

man

tle

IR2 (Olivine)IR1-o (Olivine with Cr-Mgt lamellae)

Fig. 11. Primitive mantle (McDonough & Sun, 1995) normalizedtrace element patterns of olivine in dunite (IR2) and olivinewith Cr-magnetite lamellae in orthopyroxenite (IR1-o).



Opx1V-rich vein can be explained by a monotonous in-

crease in mSiO2 and mAl2O3 from the wall rock (dunite)to the vein center. It should be noted that the

TlcþChlþTr assemblage stabilizes at higher mSiO2

and mAl2O3 conditions.

Formation conditions of Opx2Opx2 coexists with Al-rich antigorite; this assemblage is

diagnostic of eclogite-facies conditions (Bose & Ganguly,1995; Padron-Navarta et al., 2010, 2013). The texture and

compositions of minerals in IR1-o suggest a pseudo-

morphic replacement of Opx1L rims by the assemblage

Opx2þAtg 6 Chl 6 Tr6 Ilm (Fig. 3i). This textural re-

placement can be described by the balanced reaction

Opx1Lþ 0 � 0027H2O ¼ 0 � 78Opx2þ 0 � 0065Atg

þ0 � 013Chlþ 0 � 014Tr:

Whole-rock scale recrystallization to form the Opx2þAtg

assemblage is observed in localized shear zones (IR1-s)

(Figs 3e and 4f). To model the stability field of the Opx2-

bearing assemblages, P–T pseudosections were calcu-lated for the mean Opx1L rim composition in IR1-o and

the bulk composition of IR1-s in the CaO–FeO–MgO–

Al2O3–SiO2–H2O (CFMASH) system. The bulk compos-

ition of IR1-s was estimated using the volumetric propor-

tion (Atg:Opx2:Tr¼47:39:14; determined by a pixel

counting method using m-XRF maps) and the mean com-position of each mineral. Calculated diagrams were con-

toured in terms of Al contents of antigorite and

orthopyroxene. The observed Opx2-bearing assem-

blages (OpxþChlþAtgþTr in IR1-o and OpxþAtgþTr

in IR1-s) are reproduced at 1�5–2�3 GPa and 620–640�C

(Fig. 16). The very low Al content of Opx2 (<0�002

a.p.f.u.) and high Al content of antigorite (>3�0 a.p.f.u. forIR1-o and >1�2 a.p.f.u. for IR1-s) predicted in the phase-

assemblage fields are also consistent with the observed

mineral compositions in these samples. The inferred

conditions of Opx2 formation are in good agreement

with the P–T conditions of the eclogite-facies meta-

morphism (M2: 1�6–1�8 GPa and �550–620�C)


10

1

100

0.01

0.1

1

10

100

RbBa

PbTh

UNb

LaCe

PrSr

NdZr

HfSm

EuTi

GdTb

DyHo

YEr

TmYb

LuSc

VCr

CoNi

IR1-oHornblendeTremolite

IR1-vTremolite

(a) (b) Amph/C1 chondriteAmph/Primitive mantle

Fig. 12. (a) Primitive mantle (McDonough & Sun, 1995) normalized trace element patterns of amphibole in orthopyroxenite (IR1-o)and thick Opx1V-rich vein (IR1-v). (b) C1 chondrite (McDonough & Sun, 1995) normalized rare earth element patterns of amphibole.

0.001

0.01

0.1

1

10

100

1000

RbBa

PbTh

UNb

LaCe

PrSr

NdZr

HfSm

EuTi

GdTb

DyHo

YEr

TmYb

LuSc

VCr

CoNi

Opx1L

MSI+Opx1LAmph+Chl in Ol grain boundary (Fig. 3f)Phl in MSI in Opx1L (Fig. 3f)Chl (Fig. 3i)

Sample/Primitive mantle

HblHblHbl

TrTrTrChlChlChl

OlOlOl

OlOlOlPhl

Opx1LOpx1L

ChlIlm

Opx1L

TrTrTr

IR1-o

ChlIR1-v

0.0001

Fig. 13. Primitive mantle (McDonough & Sun, 1995) normalized trace element patterns and BSE images of hydrous mineral-rich do-mains in IR1-o and chlorite in IR1-v.



Olivine

Orthopyroxene

Chlorite

Antigorite

Amphibole

Cr-spinel

Pentlandite

Igneous stage Amphibolite facies stage Eclogite facies stage

high Ni

Opx1L Opx1V Opx2

Hbl (fluid origin) Tr Tr

Fe-Ti-rich Chr Ftc Cr-MagTi-poor Chr

500 600 700 800 900 1250

2.0

1.5

1.0

0.5 Ultramafic domain

M1

Pre

ssur

e (G

Pa)

Temperature (oC)

Opx1LOpx1V

M2

Opx2

DuniteMafic domainMafic domainMafic domain

+Chl

+Atg

(a)

(b)

M3

Fig. 14. (a) Evolution of mineral parageneses in outcrop IR1. (b) P–T conditions of orthopyroxene-forming stages. P–T path for themafic domain of the Western Iratsu body is taken from Endo et al. (2012). Dunite formation conditions are from Tasaka et al. (2008).

-745

-740

-735

-730

-725

-890 -885 -880-895-1665

-1660

-1655

-1650

-1645

-890 -885 -880-895

Ol

ChlTlc

Opx Tlc

Chl

Tlc

Chl

Tlc

Chl

Tlc

Ol TlcOpx Tlc

Opx

Opx

Opx

Tr

Cpx

μSiO2 (kJ)

μAl 2O

3 (k

J)

μCaO

(kJ)

μSiO2 (kJ)

Ol Tr

Ol Cpx

Opx Tr

Cpx Tr

Tr Tlc

Ol O

px

Ol Chl

Ol O

px

Opx

Opx

Opx

Opx Chl

CFMSH (Mg#=90, H2O in excess)660 oC, 1.2 GPa

0.002

0.004

0.010

FMASH (Mg#=90, H2O in excess)660 oC, 1.2 GPa

Al (apfu) in Opx

(a) (b)

Fig. 15. Isothermal and isobaric (a) mCaO–mSiO2 and (b) mAl2O3–mSiO2 pseudosections for Opx1V-rich vein (IR1-v). Arrows indicatechemical potential gradients from the wall rock to the vein center.



0.7

1.1

1.5

1.9

2.3

2.7

610 670 730 790 850

Opx Ol Cpx

Opx Ol Tr

Chl Tlc Ol Tr

Opx Atg Cpx

Opx Atg Tr

Opx Atg Tr

Opx Atg Tr

IR1-s (Opx-Tr-Atg schist)

Opx Ol Cpx

Chl Tlc Ol Tr

0.7

1.1

1.5

1.9

2.3

2.7

610 670 730 790 850

Opx Opx ChlChlAtg Atg TrTr

Opx ChlAtg Tr

1.0 1.2

1.41.6

2.03.0

2.0 2.2

2.4

3.0

3.0

(b)

(a)

Pre

ssur

e (G

Pa)

Temperature (oC)

Pre

ssur

e (G

Pa)

Temperature (oC)

Opx Opx Ol Ol Atg Atg CpxCpx

Opx Ol Atg Cpx

Opx Opx Ol Ol Atg Atg TrTr

Opx Ol Atg Tr

Opx Chl Ol CpxOpx Chl Ol CpxOpx Chl Ol Cpx

Opx Chl Ol TrOpx Chl Ol TrOpx Chl Ol Tr

Tlc Ol Atg TrTlc Ol Atg TrTlc Ol Atg Tr

Opx Tlc Opx Tlc Atg CpxAtg CpxOpx Tlc Atg Cpx

Opx Tlc Opx Tlc Atg TrAtg Tr

Opx Tlc Atg Tr

Tlc Ol Tlc Ol Atg TrAtg TrTlc Ol Atg Tr

Opx Tlc Opx Tlc Atg TrAtg Tr

Opx Tlc Atg Tr

Opx Grt Opx Grt Ol CpxOl CpxOpx Grt Ol Cpx

Opx Chl Ol TrOpx Chl Ol TrOpx Chl Ol Tr

Opx Chl Ol CpxOpx Chl Ol CpxOpx Chl Ol Cpx

Opx Opx ChlChlAtg Atg CpxCpx

Opx ChlAtg Cpx

Opx Tlc Opx Tlc Atg CpxAtg CpxOpx Tlc Atg Cpx

0.00

2

0.00

4

0.01

0

0.02

0

0.03

00.

040

0.00

2

0.00

4

0.01

0

0.02

0

0.030

2.0

0.002

Al (apfu) in Atg

Al (apfu) in Opx

IR1-o (Textural replacement of Opx1L)

CFMASH (H2O in excess)

CFMASH (H2O in excess)

Fig. 16. P–T pseudosections showing the stability fields of Opx2-bearing assemblages in (a) IR1-o (bulk composition in wt %:SiO2¼55�35, Al2O3¼1�15, FeO¼7�81, MgO¼34�23, CaO¼0�73) and (b) IR1-s (bulk composition in wt %: SiO2¼54�02, Al2O3¼0�76,FeO¼6�32, MgO¼36�96, CaO¼1�94).



determined for the mafic domain of the Western Iratsu

body (Endo, 2010; Endo et al., 2012) (Fig. 14b).

DISCUSSION

Provenance of the ultramafic domainUltramafic blocks in the Sanbagawa belt are derivedfrom dunite and subordinate wehrlite and clinopyroxen-

ite (Kunugiza et al., 1986). Residual harzburgite has not

been found. Proposed origins of the dunite–wehrlite

suites in the Sanbagawa belt are (1) olivine–clinopyrox-

ene cumulates of mafic magmas (Kunugiza et al., 1986;

Tasaka et al., 2008) and (2) residual dunite and genetic-ally related cumulus wehrlite (Hattori et al., 2010). The

provenance of ultramafic rocks in the Sanbagawa belt

has been related to the shallow forearc mantle or crust–

mantle transition zone above a subducting slab (Aoya

et al., 2013a). Similar dunite–wehrlite–pyroxenite suites

in the crust–mantle transition zone are well documentedin the Kohistan Arc (Bouilhol et al., 2009) and New

Caledonia (Pirard et al., 2013). The highly depleted na-

ture (high NiO and Mg# in olivine, high Cr# and very

low TiO2 content in primary Cr-spinel) of dunite in the

Western Iratsu body is comparable with that of other

peridotite bodies in the Sanbagawa belt (e.g. Higashi-

akaishi body). This can be related to high-degree partialmelting of hot, wet, shallow mantle (i.e. boninitic mag-

matism: >1250�C at <30 km depth) (Tasaka et al., 2008),

which takes place during a period just after subduction

initiation (Stern, 2004; Ishizuka et al., 2006). Dunite–

wehrlite suites formed beneath the proto-forearc may

have migrated toward the subduction interface by man-tle convection during the early stages of the

Sanbagawa subduction zone (Fig. 17). The dunite–wehr-

lite suite is not typical for lithologies of the mantle

wedge, but could be widespread in the hanging wall of

young subduction zones.

Melt/fluid–rock interaction during Opx1LformationThe high Ni content of relict olivine (Fig. 5a) implies the

progress of an orthopyroxene- forming reaction at the

expense of olivine (Kelemen et al., 1998). Opx1L crystalsshow Ni oscillatory zoning with a core (�0�2 wt % NiO)

and a high-Ni outer zone (up to 0�4 wt % NiO) (Fig. 7c).

The NiO content in Opx1L is too high to explain

equilibrium Ni–Mg partitioning between neoblastic

orthopyroxene and relict olivine. Low-Ni halos in Opx1

(0�08–0�15 wt % NiO) are observed around olivine inclu-

sions (Fig. 3g) and the halo–inclusion pairs give(Ni/Mg)Ol/(Ni/Mg)Opx¼3�3 6 0�4, which is close to the

equilibrium values of 4�2–8�1 at 700–900�C (Podvin,

1988) and 4�5 in an olivine–orthopyroxene rock from

Cerro del Almirez, Spain (Trommsdorff et al., 1998).

Presuming depleted dunite consisting of olivine (0�4 wt

% NiO) as the protolith, complete consumption of oliv-ine to form orthopyroxene by silica addition results in

the formation of Ni-rich orthopyroxene (0�28 wt % NiO).

This value is in agreement with the mean NiO content

in Opx1L in orthopyroxenite (Fig. 7c). The initial process

in replacive Opx1L formation is dissolution of olivine

and Cr-spinel into infiltrated Si-rich hydrous melt/fluid,

forming Ni-enriched relict olivine and a fluid oversatu-rated in the orthopyroxene component (e.g. Sen &

Dunn, 1994; Rapp et al., 1999; Perchuk et al., 2013).

Subsequent crystallization of Opx1L from the fluid did

not maintain equilibrium with the relict olivine in terms

of Ni, although the oscillatory zoning with a Ni-poor

zone implies a temporal approach to equilibrium.

The protolith of Opx1L-rich rocks is thought to bedepleted dunite consisting of olivine and Cr-spinel.

Incompatible trace element composition of olivine in

IR2 may approximate the bulk composition of the dun-

ite. Accordingly, metasomatic formation of orthopyrox-

enite (IR1-o) is associated with a significant increase in

incompatible trace element abundances. The overall re-action of the dunite–melt/fluid interaction can be written

as follows:

duniteðOlþ Ti-poor ChrÞ þ SiO2-rich melt=fluid!

Opx1Lþ Fe–Ti-rich Chrþ Chlþ residual fluid:

The residual fluid should be depleted in Si (and thus

can coexist with chromite) unless the melt/rock ratio is

very high, and is thought to be an aqueous fluid be-

cause Opx1L formation temperature is too low for maficmelts. Negative crystal shapes and the constant min-

eralogy (AmpþChlþPhlþ Ilm) suggest that MSI in

Opx1L represents remnants of the residual fluid. Trace

element compositions of the residual fluid were esti-

mated using the mean trace element composition of

interstitial hornblende (probably crystallized from theresidual fluid) and amphibole/fluid partition coefficients

(DAmp–fluid). Following the approach of Marocchi et al.

(2007), a set of DAmp–fluid values (Supplementary Data

Table 2) was derived by combining experimentally

determined DCpx–fluid values (Green & Adam, 2003) and

DAmp–Cpx data from natural samples (Ionov et al., 1997;

Zack et al., 1997; Hermann et al., 2006). DAmp–fluid valuesfor Rb and Ba are taken from Zack et al. (2001), because

low abundances of these elements in clinopyroxene

lead to large errors in computed DAmp–fluid values.

Primitive mantle normalized abundances of the calcu-

lated fluid are shown in Fig. 18a, which is characterized

by positive anomalies in Rb, Ba, U and Sr, negativeanomalies in Th, Nb, Zr and Hf, and an enrichment in

LREE over HREE. These ‘subduction’ signatures are

consistent with the measured bulk composition of MSI

(þOpx1L) (Fig. 13).

Behavior of Cr-spinel during metasomatismSlab-derived hydrous melt or solute-rich fluid is rich in

Si and therefore highly reactive with mantle rocks.

According to experimental studies (e.g. Sen & Dunn,

1994; Rapp et al., 1999) and observations on veinedmantle xenoliths (Benard & Ionov, 2013), the early pro-

cess of reaction between such a Si-rich hydrous melt





and peridotite is dissolution of olivine and Cr-spinel into

the reacted melt. Euhedral Cr-spinel grains in Opx1L are

interpreted as neoblasts in equilibrium with the Si-

depleted residual fluid. The high abundance of Cr-spinel

exsolution lamellae in olivine in Opx1L-rich rocks

implies some linkage with the Opx1L-forming meta-somatism. Incorporation of Cr into relict olivine prob-

ably took place during the Cr-spinel dissolution stage,

because formation of Cr-rich olivine in equilibrium with

Cr-spinel requires extremely high temperature condi-

tions (Li et al., 1995). Decreasing Si content in the re-

sidual fluid may be linked to the precipitation of Cr-spinel neoblasts and exsolution lamellae in olivine.

The amphibolite-facies hydrous metasomatism in

the chlorite stability field led to modification of Cr-spinel

to lower Al and Mg# compositions. According to

Gervilla et al. (2012), this process can be explained by

the following reaction: 4(Mg0�7Fe0�3)CrAlO4 (Cr-spinel)þ4Mg2SiO4 (olivine)þ2SiO2aqþ 8H2O¼ 2Mg5AlSi3AlO10

(OH)8 (chlorite)þ 2(Fe0�6Mg0�4)Cr2O6 (Fe-rich Cr-spinel).

A two-step process has been proposed for the forma-

tion of ferritchromite: formation of Fe-rich Cr-spinel via

the above reaction and subsequent oxidation (Gervilla

et al., 2012). In Opx1L-bearing rocks, compositional

modification of Cr-spinel during the amphibolite-faciesmetasomatism is associated with Ti-enrichment

(Fig. 6b). Such Ti-enrichment in ferritchromite is absent

in dunite (IR2) that is least affected by the amphibolite-

facies metasomatism (Fig. 6b). Because depleted dunite

is thought to be the protolith, the source of Ti in neo-

blastic Cr-spinel/ferritchromite is probably external andis attributed to the slab-derived melt/fluid.

Fluid–rock interaction during Opx1V formationThe texture of the Opx1V–chlorite association (Fig.4e) is very similar to that reported in chlorite-harz-

burgite formed by high-pressure dehydration of

East Asia

East Asia

Sinking sla

b

(Izanagi P

late)

Asthenosphere

Intensivefluid activity

StagnantStagnantforearc mantleforearc mantleStagnantforearc mantle

Serpentinite

Serpentinite

Serpentinite

~116 Ma Proto-forearc magmatism

Refractory harzburgite

Downgoing Izanagi Plate

Mantle melting

~89 Ma

Dunite-wehrlite Dunite-wehrlite Dunite-wehrlite Ultramafic domainSub

ducti

on

interf

ace

Ductileshear zone

Eclogite Unit

Eclogite Unit

Eclogite Unit

Besshi Unit

Besshi Unit

Besshi Unit

Aqueous fluid

Hydrous melt

Mafic domain(Juxtaposition withthe ultramafic domainand detachment fromthe slab)

Opx1L

Opx1V ~660oC, 1.2 GPa

>750oC

~620oC, 1.8 GPa Time

Opx2

Fig. 17. Schematic illustrations showing the evolution of the Western Iratsu mafic–ultramafic body in the Sanbagawa subductionzone [modified from Endo et al. (2012)].



antigorite serpentinite (Padron-Navarta et al., 2011).

However, olivine in Opx1V-rich veins is a reactantphase and is present only in thin veins or the mar-

ginal zone of thick veins. Antigorite in Opx1V-rich

veins is present as a secondary phase that replaces

Opx1V and chlorite, implying partial re-equilibration

during the later eclogite-facies stage. These observa-

tions refute a deserpentinization origin for Opx1V andwe suggest that the textural similarity to the serpent-

inite-derived chlorite-harzburgite simply reflects simi-

lar formation conditions (close to the antigorite

breakdown equilibrium in the presence of aqueous

fluid).

Field and petrographic data for Opx1V-rich veins

suggest that focused fluid flow along brittle fractures is

a viable mechanism for long-distance (at least several

meters) transport of highly reactive crustal fluids into

the unserpentinized mantle (Fig. 17). The zoned struc-ture of thick metasomatic reaction veins implies limited

chemical diffusion and reaction between a channelized

flux of crustal fluids and host dunite. Moreover, low

Opx/fluid or Chl/fluid partition coefficients for incompat-

ible trace elements suggest that the metasomatic for-

mation of Opx–Chl rock along the fluid conduits

enhances advective transport of fluid-soluble traceelements into the base (i.e. chlorite stable zone beneath

the partially molten region; Grove et al., 2006) of the

mantle wedge.

Trace element characteristics of the fluid in equilib-

rium with the central zone of a thick Opx1V-rich vein

were evaluated using the mean trace element compos-ition of tremolite (Fig. 12a) and the amphibole/fluid par-

tition coefficients (Supplementary Data Table 2).

Because of lower compatibilities of most trace elements

in tremolite compared with hornblende, the experimen-

tally determined DTr–fluid values of Fabbrizio et al. (2013)

were also used. PM-normalized trace element abun-dances of the fluid calculated using the two sets of parti-

tion coefficients show similar patterns characterized by

a positive Sr spike and negative Nb and Zr spikes

(Fig. 18a). The inferred trace element characteristics of

the fluid are consistent with an origin from the slab-

derived mafic domain in the Western Iratsu body.

Negative Eu anomalies in Opx1V and tremolite sug-gest crystallization of plagioclase from the vein-forming

fluid before interaction with dunite. This is supported by

the common occurrence of albite–quartz–phengite rock

in the mafic–ultramafic transition zone (Supplementary

Data Fig. 1).

Comparison with previous studies on supra-subduction zone mantle metasomatismTalc–chlorite–tremolite rock is a common ‘hybrid’ rock

developed between serpentinite and Si-rich crustal

rocks in many oceanic subduction-type orogens. This

rock type is stable to depths corresponding to �800�Cand is considered as an important volatile and trace

element reservoir in subduction zones (Spandler et al.

2008; Marschall & Schumacher, 2012). Orthopyroxene-

rich rock is stable at >620�C and both talc- and ortho-

pyroxene-rich rocks can be present in the deep

forearc and subarc slab–mantle wedge interface. Talc–

chlorite–tremolite rock is also very common in theSanbagawa belt (Maekawa et al., 2004), but it never

occurs in direct contact with peridotite. Compared with

talc–chlorite–tremolite rock, orthopyroxene-rich rock is

stable under lower mSiO2 conditions (e.g. Fig. 15) and

develops adjacent to peridotite. Thus, orthopyroxene-

rich rock from the exhumed slab–mantle wedge sectioncan provide information on interaction between

evolved slab fluids (which have experienced complex

Fluid in equilibrium with Hbl (IR1-o)

Calculated with DTr-fluid (Fabbrizio et al. 2013)

Fluid in equilibrium with Tr (IR1-v)

0.001

0.01

0.1

1

10

100

Li Rb Ba Th U Nb La Ce Sr Zr Hf Sm Ho Y

Flui

d/P

rimiti

ve m

antle

1000

0.1

1

10

100

1000

BaPb

ThU

NbLa

CePr

SrNd

ZrHf

SmEu

TiGd

TbDy

YEr

YbLu

0.01

0.1

1

10

100

Ulten HP Hbl

Ulten UHP Hbl

Hbl in IR1-o

Am

ph/P

rimiti

ve m

antle

A

mph

/Prim

itive

man

tle Hbl in IR1-o

Tr in IR1-v

Avacha #227 vein

Avacha ‘type 1A’ vein

Lihir

(a)

(b)

(c)

BaPb

ThU

NbLa

CePr

SrNd

ZrHf

SmEu

TiGd

TbDy

YEr

YbLu

Fig. 18. (a) Primitive mantle (McDonough & Sun, 1995) normal-ized trace element abundances of fluids in equilibrium withhornblende in IR1-o and tremolite in IR1-v. (b) Trace elementabundances in hornblende in IR1-o compared with metasom-atic hornblende in chlorite peridotites (HP Hbl; Marocchi et al.,2007) and garnet peridotites (UHP Hbl; Scambelluri et al., 2006)from the Ulten zone, Eastern Alps. (c) Trace element abun-dances in hornblende in IR1-o and tremolite in IR1-v comparedwith metasomatic amphibole in orthopyroxene-rich veins fromsubarc mantle xenoliths. Data for amphibole in mantle xeno-liths are from the Avacha volcano, Kamchatka (Ishimaru et al.,2007; Benard & Ionov, 2013) and Lihir, Papua New Guinea(Gregoire et al., 2001).






fractionation processes during migration through mul-

tiple hybrid rock layers) and the mantle wedge.

There is a strong similarity in texture between

Opx1L-rich rocks and garnet orthopyroxenite from the

Maowu ultramafic massif in the Dabieshan UHP orogen(Malaspina et al., 2006). For example, both these rocks

consist mainly of coarse-grained Ni-rich orthopyroxene

that encloses rounded relict olivine grains. Malaspina

et al. (2006) proposed that the Maowu orthopyroxenite

was derived from harzburgite that had been metasom-

atized at UHP conditions (4�0 6 1�0 GPa, 750 6 50�C) by a

hydrous melt (or supercritical fluid) sourced from asso-ciated crustal rocks (felsic gneiss). We suggest that

Opx1L-rich rocks in the Western Iratsu body were

formed by a similar process to the Maowu orthopyrox-

enite, but under much lower pressure conditions. It has

been suggested that orthopyroxene-rich rocks formed

at the subarc slab–mantle interface act as a ‘filter’ ofsome components in the slab-derived metasomatic

agent, leaving a large ion lithophile element (LILE)- and

LREE-rich residual fluid that ultimately migrates to the

source region of arc magmas (Malaspina et al., 2006;

Scambelluri et al., 2006). The present study shows that

an analogous process works at shallow forearc levels ofyoung subduction zones. The Ulten zone (Eastern Alps)

is another well-studied example of metasomatism at

the subduction interface from the amphibolite facies

(Marocchi et al., 2007) to UHP conditions (Scambelluri

et al., 2006). The P–T conditions of the amphibolite-fa-

cies metasomatism of the Ulten zone are almost identi-

cal to those of Opx1L formation in the Western Iratsubody. The Ulten peridotites and Opx1L-rich rock (IR1-o)

share common characteristics in the trace element pat-

terns of metasomatic hornblende, such as negative

HFSE (Nb, Zr, Hf and Ti) anomalies and LREE to MREE

enrichment over HREE (Fig. 18b). However, hornblende

in IR1-o shows significantly lower Ba, Th and U abun-dances (Fig. 18b), probably reflecting the different

source of the metasomatic agent between the Western

Iratsu body (oceanic crustal rocks) and the Ulten zone

(continental crustal rocks).

Trace element patterns of amphibole (tremolite,

magnesiohornblende or edenite) in the Western Iratsu

orthopyroxene-rich rocks and in subarc mantle xeno-liths from oceanic subduction settings can also be com-

pared (Fig. 18c). Given that the P–T dependences of the

DAmp–fluid partition coefficients are small, this compari-

son provides insights into the compositional evolution

of metasomatic agents from the subduction interface to

the subarc mantle. Amphibole in a fibrous orthopyrox-ene-rich vein from Lihir (Papua New Guinea; Gregoire

et al., 2001) has high trace element abundances with

negative Ba and HFSE anomalies. These features are

comparable with those of hornblende in IR1-o, although

the Lihir amphibole shows a flat HREE pattern.

Amphibole in orthopyroxene-rich veins from Avacha

(Kamchatka) generally shows relative enrichment of Baover Pb, and flat to U-shaped REE patterns (Ishimaru

et al., 2007; Ishimaru & Arai, 2008; Halama et al., 2009;

Benard & Ionov, 2013). Except for the ‘type 1A (rapidly

crystallized)’ veins described by Benard & Ionov (2013),

amphibole in the Avacha orthopyroxene-rich veins

characteristically shows positive Zr–Hf anomalies

(Ishimaru et al., 2007; Ishimaru & Arai, 2008; Halamaet al., 2009; Benard & Ionov, 2013). These trace element

characteristics of the Avacha amphibole are signifi-

cantly different from those of amphibole in the Western

Iratsu orthopyroxene-rich rocks. Most metasomatic

amphibole in the Avacha xenoliths is unlikely to have

been crystallized from a pristine slab-derived agent.

CONCLUSIONS

Orthopyroxene-rich rocks from the Western Iratsu body

provide insights into hydrous metasomatism and the

fluid flow regime in an evolving (i.e. progressive cooling

to antigorite stable conditions) forearc mantle wedgeimmediately above a subducted slab. The earliest meta-

somatism probably took place immediately after sub-

duction initiation, associated with relatively high

temperature conditions (�750�C) even near the subduc-

tion interface. During this period, a Si-rich hydrous melt

or solute-rich viscous fluid sourced from the subducted

oceanic crustal rocks reacted with dunitic uppermostmantle to form Opx1L-rich layers and a residual fluid,

which had a subduction-type geochemical signature.

Cooling of the subduction interface (�660�C at

1�2 GPa) by continued subduction led to slab-derived

fluids being more diluted and transferred into the man-

tle wedge by channelized flow along brittle fractures.Interaction between a Si-rich aqueous fluid derived

from crustal rocks (the mafic domain) and depleted

dunite at the subduction interface just outside the antig-

orite stability field resulted in the formation of Opx1V-

chlorite-rich metasomatic rock. This stage is analogous

to the processes at the down-dip end of stagnant fore-

arc mantle in a thermally matured subduction zone,which is the site of intensive fluid activity. Owing to the

low compatibility of most incompatible trace elements

in orthopyroxene and chlorite, the orthopyroxene–

chlorite channel has the potential to act as a conduit for

efficient transport of incompatible element-rich crustal

fluids into the base of the mantle wedge.The last stage of orthopyroxene formation/recrystal-

lization took place in the antigorite stability field during

eclogite-facies metamorphism (�620�C and 1�6–

1�8 GPa). Limited fluid flow along antigorite-rich ductile

shear zones resulted in the preservation of the older

stage records in low-strain regions.

Previous studies have shown that talc–chlorite–tremolite rock is the typical ‘hybrid’ rock in the slab–

mantle interface and is stable from forearc to subarc

depths (�800�C). The present study verifies the exist-

ence of orthopyroxene-rich rock at the slab–mantle

wedge interface at depths corresponding to >620�C.

Orthopyroxene-rich rock is stable under lower mSiO2

conditions than talc–chlorite–tremolite rock, and thus

develops immediately adjacent to peridotite.



Consideration of these two metasomatic rock types is

essential to understand the subduction interface

processes.

ACKNOWLEDGEMENTS

The first author thanks M. Aoya for discussion on the

Sanbagawa ultramafic rocks, and H. Mori for assistance

during sample preparation. Detailed comments by A.

Benard, N. Malaspina and an anonymous reviewer sig-

nificantly improved the paper. J. Hermann is thanked

for his comments and editorial handling.

SUPPLEMENTARY DATA

Supplementary data for this paper are available at

Journal of Petrology online.

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Orthopyroxene-rich Rocks from the Sanbagawa Belt (SW Japan