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;
VC The Author 2015. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: [email protected] 1113
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
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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
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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.
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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).
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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.
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(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.
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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
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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)
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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.
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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).
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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).
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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).
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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.
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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
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
(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).
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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)
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
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.
1128 Journal of Petrology, 2015, Vol. 56, No. 6
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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.
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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).
1130 Journal of Petrology, 2015, Vol. 56, No. 6
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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
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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)].
1132 Journal of Petrology, 2015, Vol. 56, No. 6
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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).
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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.
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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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