Sci. Rep. Kanazawa Univ.
Vol.50, p.1-27, 2006

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Geochemistry of clinopyroxene in peridotites
from the Nukabira complex, Kamuikotan zone,
Hokkaido, Japan : a LA-ICP-MS study
Akihiro TAMURA and Shoji ARAI
Department of Earth Sciences, Kanazawa University
Kakuma, Kanazawa 920-1192, JAPAN
Abstract : Trace-element compositions of clinopyroxene in peridotite from the
Nukabira complex were examined to mainly discuss partial melting processes involvedin their formation. The Nukabira complex in the Kamuikotan zone, Hokkaido, Japan, is
mainly composed of mantle-derived peridotites with wide compositional variations.
The lherzolite (e.g., C# of spinel= 0.18-0.34) and depleted harzburgite (e.g., Cr#= 0.40.7) are predominant, and dunite is often surrounded by them. Samples of 8 lherzolites,
3 harzburgites and 4 dunites from the complex were selected, and their clinopyroxene
geochemical characteristics were determined by using LA-ICP-MS. Chondritenormalized trace-element patterns of clinopyroxene in the lherzolites except one sample
are characterized by strong depletion of light-REE (LREE) and middle-REE (MREE)
relative to high heavy-REE (HREE) (e.g., (Nd/Yb)N(=chondrite normalized value)< 0.004, YbN= 2.76.1). One lherzolite sample, on the other hand, has high REE abundances of clinopyroxene (e.g. CeN!0.05, YbN!5.3, (Ce/Yb)N=0.02-0.06). The trace-element patterns ofclinopyroxene in harzburgites are “spoon-shaped” being high in MREE and LREE and
low in HREE abundance (e.g., (Nd/Yb)N< 0.09, YbN= 1.0-1.7). The abundances of
MREE and LREE are variable in each harzburgite sample (e.g., CeN= 0.008-0.03, NdN=
0.02-0.08). The dunites tend to be similar in clinopyroxene compositions to surrounding peridotites. The variations of lherzolite can be ascribed not only to the difference ofmelting degree but also to conditions on melting, such as with or without residual garnet prior to spinel peridotite stability field melting. Most of the lherzolites probably experienced the melting in the garnet peridotite stability field. The further melting of lherzolites should generate the harzburgites. The clinopyroxene geochemistry, however, indicates that the harzburgite was produced by different melting processes from the lherzolite formation. The melting assisted by LREE-rich hydrous fluid is more preferable toexplain the harzburgite features and origin.

1. Introduction
Peridotites from the Kamuikotan zone, northern Japan, are considered as residual
mantle material, and are characterized by wide variation in compositions as well as byhighly depletion in incompatible elements (e.g., Katoh and Nakagawa, 1986 ; Tamura etal., 1999). Katoh and Nakagawa (1986) proposed that the Kamuikotan peridotites were2

Akihiro TAMURA and Shoji ARAI

formed by melting of the peridotite at supra-subduction zone because their compositionalvariations are remarkably similar to those of peridotites from forearc region. The formation
process of depleted harzburgite and dunite in the Kamuikotan zone has been examined byseveral workers (e.g. Yamazaki and Niida, 1992 ; Koike, 1994 ; Tamura et al. 1999 ;Kubo, 2000). Makita and Arai (1997) reported highly depleted peridotites having high-Cr#(Cr/(Cr+Al) atomic ratio) chromian spinel (>0.6) as residual peridotites after hydrous melting of more fertile peridotite. Contrasting to dominance of highly depleted peridotite in thenorthern area of the zone, lherzolitic residue has been reported from the southern area (e.g.,
Katoh and Nakagawa, 1986 ; Tamura et al., 1999). The Nukabira complex is one of thelherzolite-bearing peridotite complexes in the Kamuikotan zone. The geochemical data as
well as the origin of lherzolite, however, have not been reported yet although the petrographical feature has been well described (e.g, Katoh, 1978 ; Katoh and Nakagawa, 1993 ;Tamura et al., 1999). Katoh and Nakagawa (1993), for example, reported lherzolite andwebsterite as the least depleted ultramafic rocks from the zone. In this paper, we preliminarily report trace-element compositions of clinopyroxene in the peridotites, mainly lherzolite, from the Nukabira complex. The compositions were determined using laser ablation-inductively coupled plasma mass spectrometry (LA-ICP-MS) installed at Kanazawa University in 2002. This paper is the first report of the clinopyroxene trace-element compositions from the Kamuikotan peridotites. The clinopyroxene geochemical data, especiallyrare earth element (REE) concentrations, are useful not only to represent the petrologicalfeature but also to understand the petrogenesis, such as evaluation of partial melting degreeof mantle peridotite and magma genesis (e.g., Johnson et al., 1990). The numerical modeling of melting processes using “REE pattern” was generally established for examination oforigin of lherzolitic rocks rather than depleted harzburgite. A good example is modeling of
the mantle process of abyssal peridotite related to MORB magma genesis at mid-ocean
ridge (e.g. Johnson et al., 1990 ; Hellebrand et al., 2002). The modeling is also applied tothe origin of peridotites from ophiolitic complexes (e.g., Kelemen et al., 1995 ; Batanova
et al.,1998 ; Batanova and Sobolev, 2000 ; Bizimis et al. 2000 ; Suhr and Edward, 2001 ;Barth et al., 2002, Takazawa et al., 2003). Abundant clinopyroxene geochemical data,
therefore, have been reported, and many studies dealt with the mantle processes using themas genetic indicator (e.g., Johnson et al., 1990 ; Bizimis et al., 2000). The main purpose ofthis study is to evaluate the clinopyroxene trace-element compositions combined with petrographical data of Nukabira peridotite by comparing with the studies of abyssal andophiolitic peridotites. This work is an extended study of Tamura (1997MS, 1999MS) andTamura et al. (1999), and therefore, we used their data and selected samples for LA-ICPMS analysis.
2. Geological background
The Kamuikotan zone is distributed discontinuously over 320 km from north to south
within the Sorach-Yezo belt of a Mesozoic accretionary complex, the central axial zone ofGeochemistry of clinopyroxene in Nukabira peridotite

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Figure 1. Distribution of ultramafic rocks (black) in the Kamuikotan zone, Hokkaido, Japan. Groups andnames of ultramafic complexes are from Katoh and Nakagawa (1986) (see text).

Hokkaido, northern Japan. The Sorach-Yezo belt is composed of the Sorach Group, Yezo
Supergroup, Kamuikotan zone, Idonnappu zone and Poroshiri ophiolite (Komatsu and
Sakakibara, 1992). The Kamuikotan zone consists of metamorphic rocks, ultramafic rocks,
greenstones and sedimentary rocks. The metamorphic rocks are generally suffered fromtypical high-pressure type metamorphism, and belong to the Kamuikotan metamorphic
rocks (Miyashiro, 1961). However, low-pressure metamorphic rocks occur as a member ofophiolites, such as the Horokanai ophiolite, northern part of the belt (Ishizuka, 1980). Thus
the Kamuikotan zone is characterized by coexistence of the two types of metamorphic
rocks (Ishizuka et al., 1983). Ultramafic complexes are distributed widely throughout the
Kamuikotan zone (Figure 1), and are in fault contact with surrounding Kamuikotan metamorphic rocks, sedimentary and mafic volcanic rocks of the Sorach group and Yezo Supergroup. Most of the complexes are mainly composed of spinel peridotite of upper mantle
origin, and they are suffered from serpentinization to various extent. Katoh and Nakagawa

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Akihiro TAMURA and Shoji ARAI(1986) classified the complexes into four regional groups, such as the northern, central,southern and southernmost groups, by differences in modal compositions of peridotites andin relative quantity of ultramafic cumulates and podiform chromitites (Figure 1). The
northern group complexes are mainly composed clinopyroxene-free harzburgite and dunite.
Ultramafic cumulates and chromitite deposits are very rare. Inushibetsu complex corresponding to the basal sequence of the Horokanai ophiolite and the Takadomari complex
consists of highly refractory peridotites (e.g., Makita and Arai, 1997 ; Tamura et al., 1999).
The central group comprises small complexes around the Kamuikotan Gorge-Biei area.
The complexes are mainly composed of harzburgite and dunite with lherzolite and a small
amount of clinopyroxenite and chromitite deposits. Some of the complexes are interpreted
as submarine sliding deposits (Maekawa, 1983). The southern group consists ofclinopyroxene-poor harzburgite and dunite. Clinopyroxenite dykes and podiform-type
chromitite deposits are not uncommon. The Iwanai-dake complex located in the central
portion of the Saru-gawa complex are composed of peridotite exceptionally free from serpentinization (e.g., Arai, 1978 ; Katoh, 1978). The southernmost group is made up mainly
by clinopyroxene-bearing harzburgite and dunite with lherzolite. Ultramafic cumulates(mainly olivine-clinopyroxenite) and podiform chromitite deposits are present. The
Nukabira complex is a representative peridotite complex in this group.
3. Petrography and mineralogy of peridotites from the Nukabira complex
The Nukabira complex is approximately 20 km× 6 km in area located in the southernpart of the Kamuikotan zone (Figure 2). It is in fault contact with the Sarugawa Formationof the Sorach Group which consists of slates, sandstones, cherts and basaltic pyroclastic
rocks (Katoh, 1978 ; Niida and Katoh, 1978), and with the Yezo Supergroup. The Undivided Hidaka (Mi-bunri-Hidaka) Supergroup (Takahashi and Suzuki, 1978), recently
called “Ganpi-yama complex” (e.g., Kawamura et al., 2001), is distributed in the central
part of the Nukabira complex (Figure 2). Although the Nukabira complex is emplaced atneighbor of the Saru-gawa complex belonging to the southern group of Katoh and Nakagawa (1986) (Figures 1 and 2), they are not continuous because the structures as well as
rock types are different between the two complexes (Katoh, 1978).
The Nukabira complex is composed of spinel-lherzolite, harzburgite and dunite.
Podiform-type chromitite is relatively frequent throughout the complex. Modal ratios ofclinopyroxene and orthopyroxene are usually up to 10 vol% and 30 vol%, respectively ,although they are quite variable at the heterogeneous zone associated with dunite (Figure 3).
Dunite is various in size and shape within lherzolite and harzburgite. Pyroxene concentrations are sometimes observed at the boundary between the dunite and harzburgite/lherzolite.
Peridotites show typical protogranular and porphyroclastic textures under the microscope.
Most of olivines exhibit both wavy extinction and kink bands. Orthopyroxene and clinopyroxene range from <1 to 5 mm in diameter and sometimes bent or deformed. Anhedral tovermicular-shaped spinel is generally interstitial to olivine and pyroxne in lherzolite andGeochemistry of clinopyroxene in Nukabira peridotite

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Figure 2. Geological map of the Saru-gawa area. Localities of LA-ICP-MS samples are shown on the
Nukabira complex and Saru-gawa complex (modified after Hashimoto et al., 1961 ; Sako andOsanai, 1962 ; Takahashi and Suzuki, 1978, 1986). Note that the Nukabira and Saru-gawacomplexes belong to the southernmost and southern groups, respectively, of Katoh and Nakagawa (1986).

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Akihiro TAMURA and Shoji ARAI

Figure 3. Modal compositions of peridotites from the Nukabira complex, in terms of olivine, orthopyroxene (Opx) and clinopyroxene (Cpx). “Dun-Host” samples are peridotites collected from inhomogeneous sections adjacent to dunite (see Tamura et al., 1999). Broken line indicates discrimination between lherzolite and harzburgite (Arai, 1984).

Figure 4. Chemical compositions of spinel and olivine in peridotites from the Nukabira complex. Compositions of the LA-ICP-MS samples are highlighted. (a) Fo content of olivine vs Cr#(=Cr/(Cr+Al) atomic ratio) of spinel. OSMA : olivine-spinel mantle array, a spinel peridotite restite
trend (Arai, 1987, 1994). (b) Mg# (=Mg/(Mg+Fe2+) atomic ratio) vs Cr# of spinel. Abyssal
peridotite field from Dick and Bullen (1984) and Arai (1994). Forearc peridotite field from
Bloomer and Hawkins (1983), Bloomer and Fisher (1987), Ishii et al. (1992) and Parkinson
and Pearce (1998). Mineral compositions shown in light color are from Tamura et al. (1999).

Geochemistry of clinopyroxene in Nukabira peridotite

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harzburgite. On the other hand, spinel is euhedral to subhedral shape in dunite.
The forsterite (Fo) content of olivine in lherzolite and harzburgite is relatively constant around 91 but vary from 89.9 to 92.3. Olivine in dunite has a wider range of the Focontent from 89.5 to 93.0 (Figure 4a). The relationship between Mg# (= Mg/(Mg+Fe2+)
atomic ratio) and Cr# (= Cr/(Cr+Al) atomic ratio) of spinel is shown in Figure 4b. The Cr#shows a weak negative correlation with the Mg#. The Cr# and Mg# of spinel range from
0.18 to 0.53 and 0.54 to 0.73 in lherzolite, and range from 0.23 to 0.67 and 0.47 to 0.67 in
harzburgite, respectively (Figure 4b). The Cr# is generally higher in harzburgite than inlherzolite with some overlapping. The compositions of spinel in dunite are divided into twogroups. One shows a lower and limited range of Cr#, around 0.4, with a range of Mg# from
0.52 to 0.67. The other group has a higher Cr# range of spinel, varying from 0.64 to 0.86,
with a range of Mg# from 0.41 to 0.55. TiO2 content of spinel is very low, less than 0.5 wt% in all peridotites. The relationship between the Fo content of olivine and Cr# of coexisting spinel in peridotite is demonstrated in Figure 4a. They have a weak positive correlation,and entirely plot along the OSMA (Olivine-Spinel Mantle Array) defined as the mantle
peridotite residue trend by Arai (1987, 1994). Mg# (=Mg/(Mg+Fe) atomic ratio) of orthopyroxene varies from 0.90 to 0.92. Al2O3 and Cr2O3 contents range from 1.1 to 4.8 wt%
and from 0.5 to 1.0 wt%, respectively. The Al2O3 content is generally higher in lherzolite(2.3-4.8 wt%) than in harzburgite (1.1-3.4 wt%). The CaO contents of orthopyroxene varyfrom 0.4 to 3.7 wt%, and are mostly lower than 2.0 wt%. The Mg# of clinopyroxene varies
from 0.92 to 0.94. The Al2O3 and Cr2O3 contents vary from 1.6 to 4.8 wt% and 0.9 to 1.5
wt%, respectively. Clinopyroxene is low in TiO2 (<0.4 wt%) and Na2O (<0.3 wt%) contents. Clinopyroxene in dunite has similar ranges of the Al2O3 (1.7-4.3 wt%) and Cr2O3 (0.6
-1.5 wt%) contents to those in lherzolite and harzburgite.
4. Samples for LA-ICP-MS study
We performed LA-ICP-MS analysis of clinopyroxene for 8 lherzolites (KNS25, KNS
40T, NUK#5, NB28-1, NDP2II, NDP2V, NDP6H and NDP11A), 3 harzburgites (NDP7A,
NDP7X and NDP10H) and 4 dunites (NDP6X, NDP7C, NDP7D and KNS45T). Localityof samples is shown in Figure 2. Modal compositions and major element compositions ofsamples are highlighted in Figure 3 and Figure 4, respectively. The lherzolite samples contain 3-10 vol% of clinopyroxene and 18-30 vol% of orthopyroxene. Modal ratio of clinopyroxene in the harzburgite and dunite is less than 2 vol%. Although the sample NDP10H
contains high amount of pyroxenes (38 vol% of opx and 7 vol% of cpx), we discriminated
NDP10H as harzburgite, based on the difference in mineral compositions, such as Cr# of
spinel, from other lherzolites (Figure 4). Samples of NDP2II and NDP2V were collectedfrom lherzolites adjacent to thin dunite and pyroxenite layers (Plate I-a). NB28-1 lherzolite
may correspond to surrounding peridotite of harzburgite-dunite suite (see Figure 10b of
Tamura et al. 1999). Sample NDP6X was collected from small irregular-shaped dunite
within NDP6H lherzolite (Plate I-b). Samples of NDP7 were collected from dunite layer

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Akihiro TAMURA and Shoji ARAI

(NDP7C and NDP7D) and host harzburgite (NDP7X and NDP7A) (Plate I-c). KNS45T
was a small block totally composed of dunite, and thus lithological relation with host rockis unclear. Measured clinopyroxenes in lherzolite and harzburgite are subhedral to interstitial grains or porphyroclasts within serpentinized olivine (Plate II). Some of them areclosely related to spinel and orthopyroxene. Clinopyroxenes in a symplectic cluster withspinel and orthopyroxene were also measured (Plate II-d). Most of clinopyroxenes are associated with spinel in dunite (Plate II-g, h).
5. Analytical method
Major element analysis of minerals was carried out with JEOL JXA-8800 electronprobe microanalyzer (EPMA) at Center for Cooperative Research of Kanazawa University.
The analysis was performed under acceleration voltage of 20 kV, a specimen current of 20
nA and a beam diameter of 3 μm. The data were corrected using a ZAF program.
Trace-element (Ti, Sr, Y, Zr, Nb and REE : La, Ce, Nd, Sm, Eu, Gd, Dy, Er, Yb andLu) concentrations of clinopyroxene in peridotite were determined by a laser ablation (193nm ArF excimer : MicroLas GeoLas Q-plus)-inductively coupled plasma mass spectrome-

Figure 5. Chondrite-normalized values of detection limit for selected elements in LA-ICP-MS analysis
with different laser diameters (50, 60 and 80 μm). Also shown are detection limits reported by
Ishida et al. (2004) for comparison. See Table 1 for values in ppm. Chondrite values from Sun
and McDonough (1989).

Geochemistry of clinopyroxene in Nukabira peridotite

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Table 1. Detection limits (ppm) for measured elements (except Si and Ca) by LA-ICP-MS analysis withdifferent laser diameters. Analytical condition is 10 Hz of repetition rate, 6 J/cm2 of pulse energy and 18 elements of simultaneous measurement.

try (Agilent 7500S) (LA-ICP-MS) at the Incubation Business Laboratory Center of Kanazawa University (Ishida et al., 2004). Analytical details and quality of data are shown ina separate paper (Morishita et al., 2005). Each analysis was performed by ablating spots of
50, 60 or 80 μm in diameter for clinopyroxenes at 10 Hz with energy density of 6 J/cm2 perpulse. Signal integration times were 40 seconds for a gas background interval and 30 seconds for an ablation interval. The NIST SRM 612 glass was used as the primary calibration
standard and was analyzed at the beginning of each batch of < 3-5 unknowns, with a linear
drift correction applied between each calibration. The element concentration of NIST SRM
612 for the calibration is selected from the preferred values of Pearce et al. (1997). Data reduction was facilitated using 29Si as internal standards for clinopyroxene, based on SiO2
contents determined by EPMA, following a protocol essentially identical to that outlined
by Longerich et al. (1996). The accuracy of measurements estimated from reproducibilityof reference material (NIST SRM 614) is better than 4% in relative standard deviation for
all elements. Typical detection limit of each element for different laser diameters is listed
in Table 1 and shown in Figure 5.
6. Results
Major-and trace-element compositions of clinopyroxene and coexisting spinel and olivine are listed in Table 2. Studied lherzolite samples have low Cr# of spinel ranging from
0.17 to 0.36 within the OSMA (Figure 4). NDP6X dunite is as low in Cr# as surroundingTable 2. Major-and trace-element compositions of clinopyroxene, and chemistries of coexisting minerals (spinel and olivine)． -：below detection limit.
Mg#=Mg/(Mg+Fe) ; Cr#=Cr/(Cr+Al) ; R3+=Cr+Al+Fe3+ ; Fet=Fe2++Fe3+ ; Fe2+ and Fe3+ of spinel are calculated assuming ideal stoichiometry,
and Mg#=Mg/(Mg+Fe2+) for spinel.

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Akihiro TAMURA and Shoji ARAI

Geochemistry of clinopyroxene in Nukabira peridotite

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Akihiro TAMURA and Shoji ARAI

Geochemistry of clinopyroxene in Nukabira peridotite

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Figure 6. Variations of clinopyroxene compositions measured in this work. Grains for LA-ICP-MS
analysis are highlighted. Partial melting and refertilization trends are from Hellebrand et al.
(2002). (a) Al2O3 content vs Mg# (=Mg/(Mg+Fe) atomic ratio). Light-colored dots from Tamura (1999MS). (b) Na2O content vs TiO2 content. TiO2 contents were determined using
EPMA.

NDP6H lherzolite (Cr#=0.35). NDP7 harzburgite and dunite have high Cr# spinel increasing from 0.57 in harzburgite (NDP7X) to 0.72 in dunite (NDP7D). NDP10H harzburgite
with high modal pyroxenes also contains higher Cr#-spinels (Cr#=0.51) comparing withlherzolite. Relationship between Mg# and Al2O3 content of clinopyroxene indicates a negative correlation trend in lherzolite and harzburgite (Figure 6a). Na2O content is decreasingfrom 0.4 wt% to 0.05 wt% with decreasing of TiO2 content from 0.2 wt% to <0.05 wt% in
lherzolite and NDP10H harzburgite (Figure 6b). However, NDP7 harzburgite and dunite
contain clinopyroxene with high Na2O content (0.3-0.5 wt%) and low TiO2 content (<0.05wt%) (Figure 6b). The dunites in lherzolite or harzburgite hosts are generally similar incompositions to their host rocks except KNS45T dunite having very high Na2O content (!
0.55 wt%) of clinopyroxene and low Fo content of olivine (Fo89.5) (Figures 4 and 6b).
Trace-element compositions of clinopyroxene in lherzolites except NDP11A are characterized by strongly depletion in middle-REE (MREE) to light-REE (LREE) relative tohigh abundances of heavy-REE (HREE) ; for example, (Nd/Yb)N(=chondrite normalized value) is lessthan 0.004 and YbN ranges from 2.7 to 6.1 (Figure 7a). MREE and LREE abundances must
be quite low because they were below detection limits of the analysis. Clinopyroxene of
NDP11A lherzolite is relatively rich in REE (e.g. CeN!0.05, YbN!5.3 and (Nd/Yb)N!0.080.09). The trace-element patterns of clinopyroxene in harzburgites are “spoon-shaped”
with gentle slopes from HREE to LREE. Low abundance of HREE (e.g., Yb=1.0-1.7) andSr positive anomaly are also recognized (Figure 7b). However, NDP10H harzburgite (e.g.,
CeN!0.008, NdN!0.02) is lower in LREE and MREE abundance of clinopyroxene than
NDP7 harzburgites (e.g., CeN!0.03, NdN!0.08). NDP7 harzburgite has high LREE/HREE
ratios of clinopyroxene, for example, (Ce/Yb)N ranging from 0.02 to 0.06, with negative Ti
anomaly in trace-element pattern (Figure 7b). NDP7 dunite is similar in the pattern of14

Akihiro TAMURA and Shoji ARAI

Figure 7. Representative chondrite-normalized trace-element patterns of clinopyroxenes in peridotites
from the Nukabira complex. (a) Clinopyroxenes from lherzolite. (b) Clinopyroxenes from harzburgite (NDP7X and NDP10H) and dunite (NDP7D and KNS45T). Plots in light color indicate the data below the detection limits. Chondrite values from Sun and McDonough (1989).

clinopyroxene to NDP7 harzburgite although there are slight differences in trace-element
abundances between them. NDP6X dunite is also similar in HREE and MREE abundancesof clinopyroxene to NDP6H of host lherzolite, and the dunite clinopyroxene has low Ceabundance detected (CeN<0.008) and weak Sr anomaly (Figure 7). There is a good correlation between the Cr# of spinel and Yb concentrations of clinopyroxene in lherzolite andharzburgite, as suggested by Hellebrand et al. (2001) (Figure 8). Dunites are also plotted on
the trend of associated peridotites. However, KNS45T dunite is slightly off the trend as
well as different in trace-element pattern from others, having weak Ti positive anomaly ofclinopyroxene (Figures 7b and 8). This is possibly due to genetic differences of dunite.
Please note, however, that compositions were obtained from clinopyroxenes intergrownwith spinel (e.g., Plate II-h) in KNS45T dunite. Clinopyroxenes forming symplectic clusters with orthopyroxene and spinel (e.g., Plate I-d) have no remarkable differences in composition from other clinopyroxenes in each lherzolite.
7. Discussion
Evaluation of the clinopyroxene geochemistry
The mineral compositions of lherzolite and harzburgite from the Nukabira complex,
for example, olivine-spinel compositions plotted along OSMA (Figure 4a), indicate thatthe peridotites are restites after various degrees of partial melting under the upper mantle
conditions. The Cr# of spinel is one of the best indicators for the partial melting degrees ofthe mantle peridotite and the discrimination of the derived tectonic setting (e.g., Dick andBullen, 1984 ; Arai, 1994 ; Hellebrand et al., 2001). Hellebrand et al. (2001) proposedthat the degree of partial melting (F) can be quantitatively estimated from the Cr# of spinel
by using equation as F=10ln(Cr#spinel)+24 Cr# for Cr# between 0.1 and 0.6 (Figure 8).

Geochemistry of clinopyroxene in Nukabira peridotite

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Figure 8. Cr# of spinel vs Yb (in ppm) of clinopyroxene in Nukabira peridotite. “Melting Trend” withmelting degree defined by F=10ln(Cr#spinel)+24 (Hellebrand et al., 2001). Abyssal peridotite
field from Johnson et al. (1990), Johnson and Dick (1992), Dick and Natland (1996), Ross andElthon (1997) and Hellebrand et al. (2001, 2002).

The degree of melting for lherzolites from the Nukabira complex ranges from 7% (KNS25)
to 14% (NDP2). Hellebrand et al. (2001) also suggested that the Cr# of spinel shows goodcorrelations with Yb abundance of clinopyroxene (Figure 8). However, two lherzolites
have similar Yb abundances (=0.9-1.0 ppm) although Cr# of spinel in one lherzolite (NDP
11A) (Cr#=0.26) is higher than that of the other one (KNS25) (Cr#=0.17). The partial
melting degree can be also numerically obtained using melting equation and distribution
coefficient of trace elements between melt and residual phases (e.g., Gast. 1968 ; Shaw,
1970 ; Johnson et al., 1990). For relationship between Dy and Ti concentrations of clinopyroxenes, the degree of melting is no differences between NDP11A and KNS25 lherzolites (Figure 9). Zr and Ti concentrations of clinopyroxene are also well established forevaluation of the melting process (e.g., Johnson et al., 1990 ; Bizimis et al, 2000) (Figure10). The melting degree estimated from the trace-element compositions of clinopyroxene issystematically higher than that obtained from Cr# of spinel in the lherzolites except NDP11
A (Figures 8, 9 and 10). The melting trends shown in Figures 9 and 10 were calculated using parameters for spinel peridotite stability field (Bizimis et al., 2000). The trace-element
fractionation between melt and residual phases is mainly controlled by distribution coefficient of coexisting phases, and melting mode, for example, batch or fractional melting.
Johnson et al. (1990) revealed that the variations of the REE patterns of clinopyroxene inabyssal peridotite were partly dependent on the garnet peridotite stability fields melting aswell as the spinel peridotite stability field melting. The discrepancy between the melting
degrees estimated by Cr# of spinel and by trace-element abundances of clinopyroxene in
the lherzolites is probably due to difference of pressure conditions on melting. The Cr# of
spinel should not effectively increase during melting in the garnet peridotite stability field16

Akihiro TAMURA and Shoji ARAI

Figure 9. Dy vs Ti (in ppm) of clinopyroxene in Nukabira peridotite. Ti content determined using LAICP-MS. Partial melting trend is from Bizimis et al. (2000) and Barth et al. (2003), calculatedusing parameter of MORB source mantle as the starting material and melting mode of dry incremental batch melting at 0.1% increments. Abyssal peridotite data source as in Figure 8.
Forearc peridotite field from Parkinson et al. (1992) and Bizimis et al. (2000).

Figure 10. Ti vs Zr (in ppm) of clinopyroxene in Nukabira peridotite. Ti content determined by LA-ICPMS. Partial melting and melt percolation trends are from Bizimis et al. (2000) and Barth et al.
(2003) : Dry melting trend as in Figure 9. Refertilization-hydrous melting trend is given byhydrous melting with contribution of two end-member fluid compositions (“max” and “min”).
Melt percolation trend is given by mantle/melt interaction process suggested by Kelemen et al.
(1995). Abyssal peridotite and forearc peridotite data source as in Figure 9.

Geochemistry of clinopyroxene in Nukabira peridotite

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Figure 11. Comparison in chondrite-normalized clinopyroxene trace-element patterns between Nukabiraperidotite and other peridotites. Chondrite values from Sun and McDonough (1989). (a) Abyssal peridotite from slow and ultra-slow spreading ridges (Mid Atlantic, Southwest and Central
Indian, and America-Antarctic ridges) from Johnson et al. (1990), Johnson and Dick (1992),
Ross and Elthon (1997) and Hellebrand et al. (2001, 2002). Representative patterns (lightcolor) are shown for Vulcan Fracture Zone (FZ) (broken lines), Atlantis II FZ and Bouvet FZ(bold lines) and Marie Celeste FZ (thin lines). (b) Peridotite from the mantle sequence in the
Oman ophiolite. Type I and II basal lherzolites (lines) from Takazawa et al. (2003). Harzburgite field from Kelemen et al. (1995). (c) Abyssal peridotite from a fast spreading ridge (Hess
Deep, East Pacific Rise) from Dick and Natland (1996). (d) Forearc peridotite field from Ishii
et al. (1992) and Parkinson et al. (1992).

(e.g., Hellebrand et al., 2002 ; Barth et al., 2003). The trace-element patterns of clinopyroxenes clearly reflect the differences of the melting process (Figure 7a). Several studies
constructed and discussed the modeling of the partial melting processes and genetic environments of abyssal and ophiolitic peridotites using the trace-element patterns of clinopyroxne (e.g., Johnson et al., 1990 ; Batanova and Sobolev, 2000 ; Suhr and Edward, 2001 ;Hellebrand et al, 2002). The moderately LREE-depleted patterns of abyssal peridotite, forexample, from the Vulcan Fracture Zone, can be reproduced by fractional melting of
MORB source mantle in the spinel peridotite stability field alone (Johnson et al., 1990)

18

Akihiro TAMURA and Shoji ARAI

(Figure 11a). On the other hand, reproduction of the clinopyroxene trace-element patterns
with the strong depletion of MREE and LREE, and with high-HREE abundances of peridotite such as those from the Bouvet and Marie Celeste fracture zones, needs melting process in the garnet-stability field prior to spinel-stability field (Johnson et al., 1990 ; Hellebrand et al., 2002) (Figure 11a). These studied patterns indicate that the different melting
histories were experienced between NDP11A and KNS25 lherzolites. NDP11A lherzolite
was probably experienced no partial melting in the garnet peridotite stability field. This is
supported by the fact that both Cr# of spinel and trace-element abundances of clinopyroxene indicate the similar melting degree (Figure 8, 9 and 10). Variations of MREE andLREE-depleted lherzolites are probably caused by melting in the spinel-stability field after
melting in the garnet-stability field because there is a good negative correlation between
the Cr# of spinel and Yb concentration of clinopyroxene in the lherzolites (Figure 8). However, the difference of melting degree in the garnet peridotite stability field can not be ruledout in some of the lherzolites. Lherzolite with ultra-LREE-depleted clinopyroxene has beenreported from the mantle sequence of the Oman ophiolite (Takazawa et al., 2003) (Figure11b). The Type I lherzolite reported by Takazawa et al. (2003) is quite similar in the traceelement pattern to the lherzolites from the Nukabira complex. Takazawa et al. (2003) proposed that the Type I lherzolite was formed by melting processes in the garnet peridotite
stability field, based on whole rock geochemistries
The degree of partial melting of the harzburgite, based on the Cr# of spinel and traceelement abundances, is simply estimated for more than 17% and up to 25% (Figures 8, 9
and 10). The clinopyroxene in NDP7 harzburgites is high in Na2O content (!0.3 wt%)
with limited low TiO2 content (< 0.05 wt%) comparing with NDP10H harzburgite clinopyroxene (Na2O< 0.1 wt %, TiO2<0.05) (Figure 6). Evidence for the high melting degree,
such as high-Cr# of spinel and low-Yb abundances of clinopyroxene, suggests that the harzburgite was probably produced by further melting of residual lherzolites. However, it isdifficult to discuss genetic relations between lherzolites and harzburgite from the Nukabiracomplex because of the discontinuity of the Cr# of spinel between studied lherzolites andharzburgites. The lherzolite variation produced by various melting environment also complicates the relations. Contrasting to the trace-element patterns of lherzolite clinopyroxenes,
enrichment in LREE and strong Sr positive anomaly are remarkably observed in the harzburgite clinopyroxenes (Figure 7b), and these features suggest that melting processes
forming harzburgite were not a series of the process forming the lherzolite variation. The
melting process should accompany refertilization observed in the clinopyroxene compositions (Figures 6b and 7b). The refertilization-hydrous melting model suggested by Bizimis
et al. (2000) is one of the ways to explain the harzburgite origin, based on the compositional relationships in clinopyroxene (Figure 10). This model is clearly distinguished frommantle/melt interaction model by the compositional trend (Figure 10) (Bizimis et al., 2000).
Tamura et al. (1999) proposed that the harzburgite is a reaction product related to dunite
formation. This interpretation was suitable to explain the formation of dunite which is not aGeochemistry of clinopyroxene in Nukabira peridotite

19

simple residue because the reaction with melt of deeper origin can effectively remove pyroxenes. The harzburgite, therefore, can be formed by re-melting of lherzolite in disequilibrium with the melt (Tamura et al., 1999). Although no modal evidence strongly indicatinghydrous condition during the melting, such as presence of hydrous minerals, was observedin the harzburgite, cryptic variations shown in this study suggest that the hydrous fluid derived from subducted slab (e.g., Bizimis et al., 2000) was associated with enhanced meltingof lherzolite. Bizimis’s model was constructed in order to interpret the origin of suprasubduction zone peridotites (Bizimis et al., 2000). Peridotite from forearc region represents
the supra-subduction zone material (e.g., Pearce et al., 1992 ; Parkinson et al., 1992), andthe Nukabira harzburgite is in good agreement in clinopyroxene geochemistry as well as
spinel compositions with the forearc peridotite (Figures 4, 9, 10 and 11c). Therefore, theharzburgite was probably generated in a supra-subduction zone setting. However, abyssal
peridotite from a fast spreading ridge, Hess deep, East Pacific Rise, is exceptionally depleted, such as high-Cr# spinel (!0.6) (Dick and Natland, 1996) (Figures 4 and 8). The
trace element patterns of clinopyroxene exhibit enrichment in MREE and LREE with low
HREE as those of the harzburgite and forearc peridotites (Figure 11c,d).
8. Conclusions and geological implications
The origin of Nukabira peridotite was examined using clinopyroxene geochemistry
determined by LA-ICP-MS. The clinopyroxene in lherzolites is distinctively depleted in
MREE and LREE. Its compositional characteristics were obtained by various melting degrees and processes, including or excluding the garnet peridotite stability field melting.
The clinopyroxene tends to be enriched in MREE and LREE in the harzburgites. The harzburgite was produced by the hydrous fluid-assisted melting of lherzolite.
For the lherzolite formation, the multi-stage melting is favorably brought by decompression of the upper mantle material. This is considered as a common process to generate
MORB magma beneath the mid-ocean ridge (e.g., Kelemen et al., 1995). The abyssal peridotite that experienced melting in the garnet peridotite stability field has been reported
from hotspot-related mid-ocean ridge setting, for example the Bouvet Fracture Zone, although the signature can not restrict to that setting (Johnson et al., 1990). The Nukabiralherzolite, thus, was probably derived from the mantle beneath the mid-ocean ridge, possibly related to the hotspot. The formation of the Nukabira harzburgite, on the other hand,should have been accomplished after the transition of tectonic setting to a supra-subduction
zone. The lherzolite probably represents the portion that survived the melting in the suprasubduction zone. Katoh and Nakagawa (1992, 1993) proposed that the southernmost groupperidotites, including the Nukabira peridotite, of the Kamuikotan zone have been derived
from the deeper part of the supra-subduction zone mantle, whereas the northern group peridotites was from shallower part. Their interpretation is consistent with our result that the
Nukabira lherzolite geochemically corresponds to the basal lherzolite in the Oman
ophiolite.

20

Akihiro TAMURA and Shoji ARAIAcknowledgments
Y. Ishida is thanked for daily discussion and assistance with the LA-ICP-MS work. Y.
Shimizu and K. Tazaki are acknowledged for assistance with the electron microprobe work.
M. Makita, T. Morishita and A. Ishiwatari are acknowledged for comments and helping
field work. This study was supported by Kanazawa University 21st Century COE program“Environmental Monitoring and Prediction of Long-and Short-Term Dynamics of PanJapan Sea Area : Construction of Monitoring Network and Assessment of Human Effects”
led by Prof. K. Hayakawa. Editorial comments by T. Morishita greatly improved thismanuscript.
Reference
Arai, S. (1978) Chromian spinel lamellae in olivine from the Iwanai-dake peridotite mass, Hokkaido, Japan. Earth and Planetary Science Letters, v.39, p.267-273.
Arai, S. (1984) Igneous mineral equilibria in some Alpine-type Peridotites in Japan. In : I. Sunagawa(Ed.) Materials Science of the Earth’s Interior. Tokyo : Terra Scientific Publishing Company,
p.445-460.
Arai, S. (1987) An estimation of the least depleted spinel peridotite on the basis of olivine-spinel mantle
array. Neues Jahrbuch für Mineralogie Monatshefte, v.8, p.347-354.
Arai, S. (1994) Characterization of spinel peridotites by olivine-spinel compositional relationships : Review and interpretation. Chemical Geology, v.113, p.191-204.
Barth, M. G., Manson, P. R. D., Davies, G. R., Dijkstra, A. and Drury, M. R. (2003) Geochemistry of the
Othris ophiolite, Greece : evidence for refertilization? Journal of Petrology, v.44, p.1759-1785.
Batanova, V. G. and Sobolev, A. V. (2000) Compositional heterogeneity in subduction-related mantle
peridotites, Troodos massif, Cyprus. Geology, v.28, p.55-58.
Batanova, V. G., Suhr, G. and Sobolev, A. V. (1998) Origin of geochemical heterogeneity in the mantle
peridotite from the Bay of Islands ophiolite, Newfoundland, Canada : Ion probe study of clinopyroxene. Geochimica et Cosmochimica Acta, v.62, p.853-866.
Bizimis, M., Salters, V. J. M. and Bonatti, E. (2000) Trace and REE content of clinopyroxenes fromsupra-subduction zone peridotite. Implications for melting and enrichment processes in island arcs.
Chemical Geology, v.165, p.67-85.
Bloomer, S. H. and Hawkins, J. W. (1983) Gabbroic and ultramafic rocks from the Mariana trench : anisland arc ophiolite. In : D. E. Hayes (Ed.) AGU Geophysical Monograph 23. The Tectonics andGeologic Evolution of Southeast Asian Seas and Islands : Part II. Washington, D. C. : American
Geophysical Union, p.294-317.
Bloomer, S. H. and Fisher, R. L. (1987) Petrology and geochemistry of igneous rocks from the Tonga
Trench-A Non-Accreting Plate Boundary. Journal of Geology, v.95, p.469-495.
Dick, H. J. B. and Natland, J. H. (1996) Late-stage melt evolution and transport in shallow mantle beneath the East Pacific Rise. In : C. Mével, K. M. Gillis, J. F. Allan et al. (Eds.) Proceedings of the
Ocean Drilling Program, Scientific Results, 147. College Station, Texas : Ocean Drilling Program,
p.103-134.

Geochemistry of clinopyroxene in Nukabira peridotite

21

Dick, H. J. B. and Bullen, T. (1984) Chromian spinel as a petrogenetic indicator in abyssal and alpinetype peridotites and spatially associated lavas. Contributions to Mineralogy and Petrology, v.86,
p.54-76.
Gast, P. W. (1968) Trace element fractionation and the origin of tholeiitic alkaline magma types. Geochimica Cosmochimica Acta, v.32, p.1057-1086.
Hashimoto, S., Suzuki, M. and Osanai, H. (1961) Explanatory text of the geological map of Japan 1 :50,000, Poroshiri-dake. Geological Survey of Hokkaido, 46p (in Japanese with English abstract).
Hellebrand, E., Snow, J. E., Dick, H. J. B. and Hofman, A. W. (2001) Coupled major and trace elements
as indicators of the extent of melting in mid-ocean-ridge peridotites. Nature, v.410, p.677-681.
Hellebrand, E., Snow, J. E., Hoppe, P. and Hofmann, A. (2002) Garnet-field melting and late-stage refertilization in ’residual’ abyssal peridotites from the Central Indian Ridge. Journal of Petrology, v.43,
p.2305-2338.
Ishida, Y., Morishita, T., Arai, S. and Shirasaka, M. (2004) Simultaneous in-situ multi-element analysisof minerals on thin section using LA-ICP-MS. The Science Reports of Kanazawa University, v.48,
p.31-42.
Ishii, T., Robinson, P. T., Maekawa, H. and Fiske, R. (1992) Petrological studies of peridotites from diapiric serpentinite seamaunts in The Izu-Ogasawara-Mariana forearc, LEG125. In : P. Fryer, J. A.
Pearce and L. B. Stokking (Eds.) Proceedings of the Ocean Drilling Program, Scientific Results,
125. College Station, Texas : Ocean Drilling Program, p.445-485.
Ishizuka, H. (1980) Geology of the Horokanai Ophiolite in the Kamuikotan Tectonic Belt, Hokkaido, Japan. Journal of the Geological Society of Japan, v.86, p.119-134 (in Japanese with English abstract).
Ishizuka, H., Imaizumi, M., Gouchi, N. and Bannno, S. (1983) The Kamuikotan zone in Hokkaido, Japan : Tectonic mixing of high-pressure and low-pressure metamorphic rocks. Journal of Metamorphic Geology, v.1, p.263-275.
Johnson, K. T. M. and Dick, H. J. B. (1992) Open system melting and temporal and spatial variation ofperidotite and basalt at the Atlantis II Fracture Zone. Journal of Geophysical Research, v.97, p.9219-9241.
Johnson, K. T. M., Dick, H. J. B. and Shimizu, N. (1990) Melting in the oceanic upper mantle : an ion
microprobe study of diopsides in abyssal peridotites. Journal of Geophysical Research, v.95,
p.2661-2678.
Katoh, T. (1978). The Sarugawa ultrabasic massif in Kamuikotan belt, central axial zone of Hokkaido.
Earth Science (Chikyu Kagaku), v.32, p.273-279 (in Japanese with English abstract).
Katoh, T. and Nakagawa, M. (1986) Tectogenesis of ultramafic rocks in the Kamuikotan tectonic belt,
Hokkaido, Japan. Monograph Association for Geological Collaboration Japan, v.31, p.119-135 (inJapanese with English abstract).
Katoh, T. and Nakagawa, M. (1992) Chromitite deposits in super-subduction ultramafics under Mesozoic Hokkaido-Sakhalin fore-arc setting, Far Eastern Asia. 29th International Geological Congress(abstract 3). Kyoto : p.722.
Katoh, T. and Nakagawa, M. (1993) Least depleted ultramafic rocks from the Kamuikotan and westernmost Hidaka serpentinite belts, Hokkaido, Japan. Prof. Jiro ISHII Memorial Volume, p.119-131 (in22

Akihiro TAMURA and Shoji ARAIJapanese).
Kawamura, M., Ueda, H., Nakagawa, M., Katoh, T. and Research Group of Hidaka Convergent Zone.
(2001) High P/T metamorphic minerals found from the “undivided Hidaka Supergroup” in the
Sorachi-Yezo Belt Hokkaido. Journal of Geological Society of Japan, v.107, p.237-240 (in Japanese with English abstract).
Kelemen, P. B., Shimizu, N. and Salters, V. J. M. (1995) Extraction of mid-ocean-ridge basalt from the
upwelling mantle by focused flow of melt in dunite channels. Nature, v.375, p.747-753.
Koike, W. (1994) Petrological studies on the Shirikoma-dake ultramafic complex in Kamuikotan Zoneextremely depleted chemical compositions. 101st Annual Meeting of the Geological Society of Japan (abstract). Sapporo : p.346 (in Japanese).
Kubo, K. (2002) Dunite formation processes in highly depleted peridotite : case study of the Iwanaidake
peridotite, Hokkaido, Japan. Journal of Petrology, v.43, p.423-448.
Komatsu, M. and Sakakibara, M. (1992) General geology of Hokkaido. In : M. Komatsu, H. Shibakusa,
S. Miyashita, H. Ishizuka, Y. Osanai and M. Sakakibara (Eds.) Subduction and Collision related
high and low P/T metamorphic belts in Hokkaido. 29th International Geological Congress Field trip
C01, p.1-61.
Longerich, H. P., Jackson, S. E. and Gunther, D. (1996) Laser ablation inductively coupled plasma mass
spectrometric transient signal data acquisition and analyte concentration calculation. Journal of
Analytical Atomic Spectrometry, v.11, p.899-904.
Makita, M. and Arai, S. (1997) Diversity of the highly refractory peridotite : Kamuikotan peridotite vs.
Papuan peridotite. Report of Grant-in-Aid of Monbusho “Significance of Greenstone in the Formation of Accretionary Complex”. v.2, p.97-106 (in Japanese).
Miyashiro, A. (1961) Evolution of metamorphic belts. Journal of Petrology, v.2, p.227-331.
Morishita, T., Ishida, Y., Arai, S. and Shirasaka, M. (2004) Determination of multiple trace element
compositions in thin (< 30 μm) layers of NIST SRM 614 and 616 using laser ablation-inductively
coupled plasma-mass spectrometry. Geostandards and Geoanalytical Research, v.29, p.107-122.
Niida, K. and Katoh, T. (1978). Ultramafic rocks in Hokkaido. Monograph Association for GeologicalCollaboration Japan, v.21, p.61-81 (in Japanese with English abstract).
Parkinson, I. J. and Pearce, J. A. (1998). Peridotites from the Izu-Bonin-Mariana forearc (ODP Leg
125) : evidence for mantle melting and melt-mantle interaction in a supra-subduction zone setting.
Journal of Petrology, v.39, p.1577-1618.
Parkinson, I. J., Pearce, J. A., Thirlwall, M. F., Johnson, K. T. M. and Ingram, G. (1992) Trace element
geochemistry of peridotites from the Izu-Bonin-Mariana forearc. In : P. Fryer, J. A. Pearce and L. B.
Stokking (Eds.) Proceedings of the Ocean Drilling Program, Scientific Results, 125. College Station, Texas : Ocean Drilling Program, p.487-506.
Pearce, J. A., van der Laan, S. R., Arculus, R. J., Murton, B. J., Ishii, T., Peate, D. W. and Parkinson, I. J.
(1992) Boninite and harzburgite from LEG 125 (Bonin-Mriana Forearc) : a case study of magma
genesis during the initial stages of subduction. In : P. Fryer, J. A. Pearce and L. B. Stokking (Eds.)
Proceedings of the Ocean Drilling Program, Scientific Results, 125. College Station, Texas : Ocean
Drilling Program, p.623-657.

Geochemistry of clinopyroxene in Nukabira peridotite

23

Pearce, N. J. G., Perkins, W. T., Westgate, J. A., Gorton, M. P., Jackson, S. E., Neal, C. R. and Chenery,S. P. (1997) A compilation of new and published major and trace element data for NIST SRM 610
and NIST SRM 612 glass reference materials. Geostandards Newsletter, v.21, p.115-144.
Ross, K. and Elthon, D. (1997) Extreame inconmpatible trace-element depletion of diopside in residual
mantle from south of the Kane F.Z. In : J. A. Karson, M. Cannat D. J. Miller et al. (Eds.) Proceedings of the Ocean Drilling Program, Scientific Results, 153. College Station, Texas : Ocean Drilling Program, p.277-284.
Sako, S. and Osanai, H. (1962) Explanatory text of the geological map of Japan 1 : 50,000, Chiroro.
Geological Survey of Hokkaido, 46p (in Japanese with English abstract).
Shaw, D. M. (1970) Trace element fractionation during anatexis. Geochimica Cosmochimica Acta, v.34,
p.237-243.
Sun, S. S. and McDonough, W. F. (1989) Chemical and isotopic systematics of oceanic basalts : implications for mantle composition and processes. In : A. D. Saunders and M. J. Norry (Eds.) Geological Society Special Publication 42. Magmatism in the Ocean Basins. London : Geological Society,p.313-345.
Takahashi, K. and Suzuki, M. (1978) Explanatory text of the geological map of Japan 1 : 50,000,
Iwachishi. Geological Survey of Hokkaido, 46p (in Japanese with English abstract).
Takahashi, K. and Suzuki, M. (1986) Explanatory text of the geological map of Japan 1 : 50,000,
Hidaka. Geological Survey of Hokkaido, 44p (in Japanese with English abstract).
Takazawa, E., Okayasu, T. and Satoh, K. (2003) Geochemistry and origin of the basal lherzolites fromthe northern Oman ophiolite (northern Fizh block). Geochemistry Geophysics Geosystems, v. 4(2),Tamura, A. (1997MS) Petrology of The Nukabira peridotite massif, Kamuikotan belt, Hokkaido. Graduate Thesis, Kanazawa University. 39p (in Japanese with English abstract).
Tamura, A. (1999MS) Petrogenesis of peridotites in the Kamuikotan belt, Hokkaido, northern Japan. Kanazawa, Master Thesis, Kanazawa University. 84p.
Tamura, A., Makita, M. and Arai, S. (1999) Petrogenesis of ultramafic rocks in the Kamuikotan belt,
Hokkaido, northern Japan. The Memoirs of the Geological Society of Japan, No.52, p.53-68 (inJapanese with English abstract).
Yamazaki, D. and Niida, K. (1992) Iwanai-dake peridotite in Kamuikotan Zone : Magma channels inhigh Mg-peridotites. 100 Annual Meeting of the Geological Society of Japan (abstract). Tokyo :p.575 (in Japanese).

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Akihiro TAMURA and Shoji ARAI

Plate I.
Modes of occurrence of peridotite in the Nukabira complex.
a. Thin dunite layer with pyroxenite seam within lherzolite (NDP2).
b. Irregular-shaped dunite within lherzolite (NDP6).
c. Layered dunite-harzburgite suite (NDP7). Highlighted the sampling spots (from top ; X,
A, B, C and D).
Plate II.
Photomicrographs of clinopyroxene in peridotite from the Nukabira complex. Abbreviations ; ol : olivine, opx : orthopyroxene, cpx : clinopyroxene and sp : spinel. Scale bar :1 mm.
a. Clinopyroxene porphyroclast and subhedral spinel in lherzolite (KNS25). Crosspolarized light.
b. Anhedral to interstitial clinopyroxene surrounded by serpentinized olivine in lherzolite(NDP11A). Cross-polarized light.
c. Aggregate of clinopyroxene and orthopyroxene in lherzolite (NUK#5). Cross-polarized
light.
d. Symplectic cluster of clinopyroxene-orthopyroxene-spinel in lherzolite (NUK#5). Notethat respective segments of orthopyroxene and clinopyroxene have the same optical orientation. Cross-polarized light.
e. Clinopyroxene between orthopyroxenes in harzburgite (NDP7X). Cross-polarized light.
f. Discrete clinopyroxene surrounded by serpentinized olivines in harzburgite (NDP7A).
Cross-polarized light.
g. Clinopyroxene closely associated with euhedral spinel in dunite (NDP7D). Planepolarized light.
h. Cluster of clinopyroxene-spinel included in olivine in dunite (KNS45T). Cross-polarized
light.

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Geochemistry of clinopyroxene in Nukabira peridotite

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