ORIGINAL PAPER
Fluid/rock interaction and mass transfer in continental subduction
zones: constraints from trace elements and isotopes (Li, B, O, Sr,
Nd, Pb) in UHP rocks from the Chinese Continental Scientific
Drilling Program, Sulu, East China
Yilin Xiao • Jochen Hoefs • Zhenhui Hou •
Klaus Simon • Zeming Zhang
Received: 23 September 2010 / Accepted: 7 March 2011 / Published online: 19 April 2011
 Springer-Verlag 2011
Abstract In order to better understand the role of fluids
during subduction and subsequent exhumation, we have
investigated whole-rock and mineral chemistry (major and
trace elements) and Li, B as well as O, Sr, Nd, Pb isotopes
on selected continuous drill-core profiles through contrast-
ing lithological boundaries from the Chinese Continental
Scientific Drilling Program (CCSD) in Sulu, China. Four
carefully selected sample sets have been chosen to inves-
tigate geochemical changes as a result of fluid mobilization
during dehydration, peak metamorphism, and exhumation
of deeply subducted continental crust. Our data reveal that
while O and Sr-Nd-Pb isotopic compositions remain more
or less unchanged, significant Li and/or B isotope fractio-
nations occur between different lithologies that are in close
contact during various metamorphic stages. Samples that
are supposed to represent prograde dehydration as indicated
by veins formed at high pressures (HP) are characterized by
element patterns of highly fluid-mobile elements in the
veins that are complementary to those of the host eclogite.
A second sample set represents a UHP metamorphic crustal
eclogite that is separated from a garnet peridotite by a thin
transitional interface. Garnet peridotite and eclogite are
characterized by a [10% difference in MgO, which, toge-
ther with the presence of abundant hydroxyl-bearing min-
erals and compositionally different clinopyroxene grains
demonstrate that both rocks have been derived from dif-
ferent sources that have been tectonically juxtaposed during
subduction, and that hydrous silicate-rich fluids have been
added from the subducting slab to the mantle. Two addi-
tional sample sets, comprising retrograde amphibolite and
relatively fresh eclogite, demonstrate that besides external
fluids, internal fluids can be responsible for the formation of
amphibolite. Li and B concentrations and isotopic compo-
sitions point to losses and isotopic fractionation during
progressive dehydration. On the other hand, fluids with
isotopically heavier Li and B are added during retrogres-
sion. On a small scale, mantle-derived rocks may be sig-
nificantly metasomatized by fluids derived from the
subducted slab. Our study indicates that during high-grade
metamorphism, Li and B may show different patterns of
enrichment and of isotopic fractionation.
Keywords Fluid/rock interaction  Elemental transfer 
Isotopic fractionation  Subduction and exhumation  Sulu
Introduction
It is well known that subduction of altered basaltic oceanic
crust and overlying sediments is characterized by signifi-
cant releases of aqueous fluids from metamorphosing slabs
(Bebout 2007; Forneris and Holloway 2003; Jarrard 2003;
Communicated by T. L. Grove.
Electronic supplementary material The online version of this
article (doi:10.1007/s00410-011-0625-4) contains supplementary
material, which is available to authorized users.
Y. Xiao (&)  Z. Hou
CAS Key Laboratory of Crust-Mantle Materials
and Environments, School of Earth and Space Sciences,
University of Science and Technology of China,
230026 Hefei, China
e-mail: ylxiao@ustc.edu.cn
Y. Xiao  J. Hoefs  K. Simon
Geowissenschaftliches Zentrum der Universita¨t Go¨ttingen,
Goldschmidtstrasse 1, 37077 Go¨ttingen, Germany
Z. Zhang
Institute of Geology, Chinese Academy of Geological Sciences,
100037 Beijing, China
123
Contrib Mineral Petrol (2011) 162:797–819
DOI 10.1007/s00410-011-0625-4
Peacock 1990). Since subducted oceanic material rarely
returns back to the surface, inferences about the fluid
inventory during subduction relies on indirect evidence,
such as the analysis of arc rocks. In contrast, deeply sub-
ducted continental crust is usually brought back to the
surface due to its lower density. However, only limited
attention has been paid so far to the role of fluids during
subduction and exhumation of continental crust (e.g.,
Zheng 2009). In comparison with hydrated oceanic crust,
subducted continental crust is drier, commonly older and
colder, less dense, and enriched in incompatible elements.
Nevertheless, subducted continental materials must have
been subjected to dehydration, releasing fluids from the
subducting slab. Substantial amounts of H2O may have
been transported to depths of more than 100–200 km by
hydrous minerals during UHP metamorphism (Poli and
Schmidt 1995; Thompson 1992). Volatiles released by
decomposition of hydrous phases may enter the structure of
anhydrous mantle minerals (Su et al. 2004; Xia et al. 2005).
Lithium (Li), boron (B) and oxygen (O) isotopes are
excellent tracers for fluid/rock interactions during down-
going and uplift processes of subducted slabs (e. g., Elliott
et al. 2006; Ionov and Seitz 2009; Leeman et al. 2004;
Magna et al. 2006; Moriguti and Nakamura 1998; Moriguti
et al. 2004; Teng et al. 2007; Tomascak 2004). Early Li and
B isotope analyses have highlighted the behavior of these
elements under hydrothermal and low-temperature condi-
tions, with the heavier isotopes (7Li and 11B) becoming
enriched in fluid phases (Chan et al. 1992; Hervig et al.
2002; Hoefs and Sywall 1997; Palmer and Swihart 1996;
Rudnick et al. 2004; Seyfried et al. 1998). Recently,
experimental work documented the partitioning and isoto-
pic fractionation of Li and B between fluids and minerals at
higher temperatures (e.g., Wunder et al. 2005, 2006; Meyer
et al. 2008). Interestingly, at high temperatures, B isotopes
vary regularly with the degree of B enrichment, whereas Li
isotopes do not correlate in such a general way (e.g., Elliott
et al. 2004; Tomascak et al. 2002). Although extremely low
d7Li values in HP eclogites have been reported, the cause
for the depletion of Li isotopes is unclear (Agostini et al.
2008; Marschall et al. 2007a; Simons et al. 2010; Zack
et al. 2003). Thus, apparently isotopes of the two elements
may behave differently at high-temperature conditions. At
mantle conditions, Li can be incorporated into magnesian
silicates while B is dominantly fluid controlled (Bebout
et al. 1993; Seyfried et al. 1998). Consequently, the Li and
B systematics may provide a record of fluid/rock and crust/
mantle interactions in deep subduction zones.
The Dabie–Sulu UHP metamorphic belt, which is the
largest known UHP terrain, is a suitable region to study the
fluid regime of subducted crustal material, because of its
abundant veins that formed during subduction (Castelli
et al. 1998; Franz et al. 2001) and exhumation (Wu et al.
2009; Zheng et al. 2007). Furthermore, variable extents of
fluid/melt metasomatism are evident at contacts between
different lithological units (Zhao et al. 2007). In this paper,
we outline the petrology, bulk-rock and mineral geo-
chemistry, and isotope systematics of four carefully
selected sample sets collected from drill cores of the
Chinese Continental Scientific Drilling program (CCSD),
representing different stages of fluid/rock interaction. We
investigate elemental and isotopic fractionations, during
(i) prograde metamorphism, (ii) peak metamorphic condi-
tions, and (iii) retrograde exhumation. Of special relevance
in our study are investigations into Li and B concentrations
and isotopic compositions.
Geological background and samples
The CCSD has been carried out in the Dabie–Sulu meta-
morphic belt at Donghai with a final depth of 5,158 m for
the main hole (Xu 2007). Orthogneiss, paragneiss, eclogite,
amphibolite (mostly retrograded from eclogite), and ultra-
mafic rocks are the main rock types that have been drilled;
rarely, schist, and quartzite occur as thin interlayers within
paragneiss and eclogite (for details, see Zhang et al. 2006).
The overall thickness of eclogites is about 1,200 m.
A garnet peridotite sequence, which is sandwiched by
normal or quartz-rich eclogite at the base and highly Fe–Ti-
enriched eclogite at the top, has been recovered from 603.2
to 683.5 m (Yang et al. 2007). High-pressure veins with
eclogite-facies mineral assemblages and retrograde veins
with amphibolite-facies mineral assemblages are com-
monly observed in the eclogite cores (e.g., Xu 2007; Zhang
et al. 2008). All rock types down to 5,158 m from the
CCSD have been subjected to UHP metamorphism, as
documented by the preservation of coesite inclusions in
zircon (Liu et al. 2007).
Samples for this study were collected from small-scale
CCSD continuous profiles (Fig. 1). After general inspection
of the drill cores, we carefully selected four sets of samples
that in our view are rocks formed during different meta-
morphic stages, ranging from (i) prograde dehydration
through (ii) metamorphic peak conditions to (iii) retro-
gression during exhumation. Sample sets of 1, 3, and 4 were
collected from 312.7 (CCSD-1-Ec and CCSD-1-Hv), 369.1
(CCSD-3-Ec and CCSD-3-Am), and 511.0 m (CCSD-4-Ec
and CCSD-4-Am) of the CCSD main hole, whereas sample
set 2 (CCSD-2-Gp, CCSD-2-Tr, and CCSD-2-Ec) was from
drill-hole ZK703, which is about 80 m away from the main
hole (Xiao et al. 2006). Sample set 1 (Fig. 1a) contains
hydrous veins and consists of high-pressure mineral
assemblages like phengite, quartz, and clinopyroxene
formed during prograde dehydration of eclogites that are
regarded to represent direct samples formed during
798 Contrib Mineral Petrol (2011) 162:797–819
123
subduction dehydration (Becker et al. 1999; Castelli et al.
1998; Spandler and Hermann 2006).
Peridotites enclosed within UHP metamorphic rocks
(sample set 2; Fig. 1b) provide an excellent natural labo-
ratory to study fluid-controlled metasomatism in a slab/
mantle system (e.g., Marocchi et al. 2007, 2009; Malaspina
et al. 2009). Of particular interest are garnet peridotites, as
a minimum pressure of 1.6 GPa is required to stabilize
garnet under lherzolitic compositions (e.g., Green and
Ringwood 1967; O’Neill 1981; Robinson and Wood 1998).
Retrograde veins (sample set 3 and 4; Fig. 1c, d) formed
during uplift and cooling typically contain amphibole and
plagioclase that differ from the phase assemblages in high-
pressure veins and thus are suitable to study fluid/rock
interaction during exhumation (Barnicoat and Fry 1986).
Mineral assemblages of investigated samples are shown
in Table 1. Metamorphic P–T conditions of the sampled
eclogites and garnet peridotite were calculated using
experimentally and empirically calibrated geothermome-
ters and geobarometers involving equilibria of garnet (Grt),
omphacite (Cpx), orthopyroxene (Opx), and phengite. The
minimum pressure of peak metamorphism for all investi-
gated eclogites can be assumed to be[2.8 GPa, because of
inclusions of quartz pseudomorphs after coesite and
because of coesite inclusions in zircons (Liu et al. 2007).
The results of P–T calculations are summarized in Table 2.
Sample set 1: dehydrated eclogite (CCSD-1-Ec)
and high-pressure vein (CCSD-1-Hv)
This sample set consists of an eclogite and a cm-thick
‘‘pegmatite-like’’ vein (Fig. 1a). The host eclogite (CCSD-
1-Ec) contains garnet, omphacite, phengite, and minor
quartz and rutile, which may represent a dehydrated relic.
Generally, minerals in the eclogite are relatively fresh and
free from visible retrogression. However, slight retrogres-
sion is evident by the presence of amphibole after
omphacites and phengites with low SiO2 in the rims. The
pegmatite-like HP ‘‘vein’’ (CCSD-1-Hv; *2 cm wide)
contains coarse-grained phengite (up to 1 cm in size,
Fig. 1a), quartz, omphacite, amphibole (at the rim of
omphacite), and rare garnet. The coarse-grained phengite is
parallel to the strike direction of the vein. Omphacite
occurs both in the matrix and as inclusions in garnet
(Fig. 2a). Omphacite in the matrix may be partially
replaced and rimmed by late symplectite of clinopyrox-
ene ? sodic amphibole ? albitic plagioclase. Rutile and
apatite are accessory phases.
Using the Grt–Cpx–phengite geobarometer of Carswell
et al. (1997) for temperatures of 800C, pressures of
2.4–2.7 GPa at the rim and of 3.4–3.5 GPa at the core were
obtained for phengites from the eclogite and from the vein
of 2.6–2.9, 3.4–3.5, and 2.5–2.6 GPa at the rim, the central
Fig. 1 Investigated samples showing small-scale distinguishable
lithologies collected from drill cores of the Chinese Continental
Scientific Drilling Program. a Shows a HP vein in the center (CCSD-
1-Hv) and dehydrated eclogite on both sides (CCSD-1-Ec);
b represents the transition from garnet peridotite to eclogite with a
transition zone; c and d comprise retrograde amphibolite and
relatively fresh eclogite
Contrib Mineral Petrol (2011) 162:797–819 799
123
portion, and the core, respectively (Table 2). Tempera-
tures, estimated by applying the garnet–clinopyroxene
geothermometers of Ellis and Green (1979); Krogh (1988)
and Aranovich and Pattison (1995) in coexisting ompha-
cite–garnet pairs, are 760–820C for mineral rims and
800–850 for mineral cores. In summary, the HP vein and
the host eclogite have been subjected to similar peak
metamorphic pressure and temperature conditions, which
are around 3.5 GPa at temperatures around 800C.
Sample set 2: transition from mantle-derived rocks
to eclogite
This sample set includes three distinct lithologies (a garnet
peridotite, a transition zone, and an eclogite), representing a
UHP metamorphic crustal rock that is separated from a
mantle-derived rock by a thin transition zone (Fig. 1b). The
garnet peridotite (CCSD-2-Gp) consists of olivine, garnet,
clinopyroxene, and ilmenite. Chromite and magnetite are
accessory phases. Olivine usually shows an interconnected
network of serpentinization (Fig. 2b). Exsolution of
ilmenite in olivine is a very common feature (Fig. 2c, d).
Orthopyroxene has not been observed. The absence of
spinel and the preservation of abundant ilmenite exsolutions
in olivine grains, while lacking titanoclinohumite, indicate a
minimum formation pressure of[5 GPa (Dobrzhinetskaya
et al. 1996). Temperatures estimated from Grt–Cpx geo-
thermometers vary from 1,000 to 1,060C.
The transition zone (CCSD-2-Tr) is about 2 cm thick
and contains high modal abundances of coarse-grained
omphacite (*60%, up to 3–5 mm in size;) and garnet
(*30%, *1 mm in size), and minor olivine and ilmenite.
Chlorite and amphibole are obviously products of fluid/
rock interaction (Figs. 2d, e, 3). Ilmenite, chromite, and
sulfide phases (FeS) are present in rare cases (Fig. 2f).
Very rare orthopyroxene (a few grains on thin-section
scale) coexists with amphibole (Fig. 3). The transition zone
represents an interface between the garnet peridotite and
the eclogite. Based on mineral assemblage, it is part of the
eclogite rather than that of the garnet peridotite. The con-
tact between the garnet peridotite and the transition zone is
very sharp, indicating no genetic relation between the
peridotite and the eclogite (see also Yang et al. 2007). The
formation of coarse-grained clinopyroxene (Cpx I) in
the transition zone (Fig. 1b) and the occurrence of sec-
ondary clinopyroxene (Cpx II, Fig. 3) indicate that eclogite
and garnet peridotite came together in close contact under
high P–T conditions.
Homogeneous coarse-grained omphacite (Cpx-I, with
Na2O [4 wt%) yields temperatures of 1,015–1,070C
applying the above-mentioned Grt–Cpx thermometers.
Application of the Grt–Opx thermometer of Lee and
Ganguly (1988) for zoned Opx yields temperatures of 750T
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800 Contrib Mineral Petrol (2011) 162:797–819
123
in the core and 790 in the rim rapidly increasing to 960C
at the outermost rim. The corresponding pressure estimates
from the Opx–Grt barometer of Harley and Green (1982)
are 2.5–2.8 GPa in the core that increase to 3.3 GPa in the
rim. The recrystallized Cpx-II with variable but usually
\3.5 wt% Na2O contents is also zoned, giving tempera-
tures around 780C in the core and increasing to 1,020C in
the rim (Fig. 4). The rimward increase in P–T estimates for
the zoned Opx and Cpx-II indicates that the two minerals
grew under prograde metamorphic conditions.
Eclogite (CCSD-2-Ec) consists of garnet (*55%),
clinopyroxene (*40%), and ilmenite (*5%), similar to
the mineral assemblages of the transition zone, but the size
of mineral grains is much smaller. No olivine, chlorite, or
amphibole has been found in the eclogite. Peak metamor-
phic temperature conditions calculated for the eclogite are
950–1,020C, in agreement with those calculated for the
garnet peridotite considering the error ranges of the used
thermometers.
Sample sets 3 and 4: samples representing retrograde
metamorphism
Sample sets 3 and 4 represent coexisting eclogite and ret-
rograde amphibolite, in which the behavior of elements and
isotopes during retrograde fluid/rock interaction can be
studied. Sample set 3 comprises a relatively fresh (little
altered) eclogite (CCSD-3-Ec) and a retrograde amphibo-
lite (CCSD-3-Am) that formed from retrogression of
eclogite during exhumation. On the hand specimen scale
(about 10 cm), the boundary between the amphibolite and
the eclogite is clearly obvious (Fig. 1c). The eclogite
(CCSD-3-Ec) consists mainly of garnet, omphacite, and a
few percent of rutile. Omphacite close to the boundary to
amphibolite shows only slight retrogression at the rim of
grains. Amphibolite (CCSD-3-Am) contains low-pressure
amphibole and plagioclase with rare garnet and ilmenite
but no omphacite. However, abundant omphacite inclu-
sions in rarely preserved garnet point to the same UHP
metamorphism as the eclogite. This altogether indicates
that fluid addition induced the retrogression of amphibolite
from the almost hydroxyl-free precursor assemblage.
Sample set 4 comprises a retrograde amphibolite-vein,
being a few centimeters wide, and a well-preserved host
eclogite (Fig. 1d). The amphibolite vein (CCSD-4-Av)
tends to be filled with symplectite (amphibole ? plagio-
clase) and usually minor quartz, the amount of which
increases with the thickness of the vein. The host eclogite
(CCSD-4-Ec) consists of omphacite ? garnet ? quartz ?
minor rutile.
Table 2 Pressure–temperature estimate of the investigate samples
Sample Temperature calculation Pressure calculation
Mineral T (C, at 2.8 GPa) Mineral P (GPa, at 800C)
CCSD-1-Ec Cpx–Grt (r) 760–820 (EG, KR, AP) Phen (r) 2.4–2.7 (CA)
Cpx–Grt(c) 800–850 (EG, KR, AP) Phen (c) 3.4–3.5 (CA)
CCSD-1-Hv Cpx(in)–Grt 780–840 (EG, KR, AP) Phen (r) 2.6–2.9 (CA)
Cpx–Grt(c) 810–840 (EG, KR, AP) Phen (m) 3.4–3.5 (CA)
Qtz–Phen 790–820 (ZH) Phen (c) 2.5–2.6 (CA)
CCSD-2-Gp Cpx–Grt (c & r) 960–1050 (EG, KR, AP) Ilmenite exsolutions in olivine [5.0 (DO)
CCSD-2-Tr Cpx(I)–Grt 960–1060 (EG, KR, AP)
Cpx(II, c)–Grt 770–890 (EG, KR, AP)
Cpx(II, r)–Grt 970–1030 (EG, KR, AP)
Opx (c) 700–790 (LG, HA) Opx (c) 2.5–2.8 (HG)
Opx (r) 940–960 (LG, HA) Opx (r) 3.4 (HG)
CCSD-2-Ec Cpx–Grt (c) 800–900 (EG, KR, AP)
Cpx–Grt (r) 910–1000 (EG, KR, AP)
CCSD-3-Ec Cpx–Grt 815–840 (EG, KR, AP) [2.8 (LI)
CCSD-3-Am Cpx(in)–Grt 800–825 (EG, KR, AP) [2.8 (LI)
Am–Grt 760 (RA)
CCSD-4-Ec Cpx–Grt 800–850 (EG, KR, AP) [2.8 (LI)
CCSD-4-Av Cpx(in)–Grt 800–850 (EG, KR, AP) [2.8 (LI)
Qtz–Grt 850 (ZH)
Abbreviations for applied thermobarometries: AP Aranovich and Pattison (1995), CA Carswell et al. (1997), DO, Dobrzhinetskaya et al. (1996),
EG Ellis and Green (1979), KR Krogh (1988), HA Harley (1984), HG Harley and Green (1982), LG Lee and Ganguly (1988), LI Liu et al. (2007),
RA Ravna (2000), ZH Zheng (1993). Other abbreviations: Am amphibole, Cpx clinopyroxene, Grt garnet, Phen phengite, Opx orthopyroxene, Qtz
quartz, c core, m middle, r rim; Cpx (in) inclusion in garnet
Contrib Mineral Petrol (2011) 162:797–819 801
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Peak metamorphic conditions for CCSD-3-Ec were
calculated to be 815–840C at 2.8 GPa using coexisting
Grt–Cpx pairs in the matrix (Table 2). Those for retrograde
amphibolite (CCSD-3-Am) were estimated to be 810–825C
based on the presence of omphacite inclusions in garnet. For
CCSD-4-Ec and CCSD-4-Av, peak metamorphic tempera-
tures are similar, ranging from 800 to 850C at pressure of
2.8 GPa.
Analytical methods
Major and trace elements
Bulk major elements were determined by X-ray fluores-
cence (XRF). Glass beads used for analysis were made by
mixing with lithium tetraborate as a flux. XRF analysis was
undertaken using a Philips PW 1480 automated sequential
Fig. 2 Backscattered electron (BSE) images of investigated samples
illustrating petrological features (see text). a Coarse-grained phengite,
garnet, and omphacite indicate the high-pressure origin of sample
CCSD-Hv; slight retrogression can be observed at the rims of phengite
and omphacite grains. b Olivine is the major mineral phase of the
garnet peridotite (CCSD-2-Gp); however, retrograded serpentine is
also visable. c Abundant ilmenite exsolutions in olivine grains refer to
very high pressure for the garnet peridotite (Dobrzhinetskaya et al.
1996). d and e Abundant hydroxyl-bearing minerals occur in the
transition zone (CCSD-2-Tr) between garnet peridotite and eclogite.
f Intergrown chromite and sulfide phases (FeS) coexisting in the
transition zone (CCSD-2-Tr)
802 Contrib Mineral Petrol (2011) 162:797–819
123
spectrometer. Trace elements of bulk rocks were obtained
by ICP-MS solution analyses. Samples were digested
with HF, HClO4, and HNO3 at 200C for a few days in
screw-top Teflon vials before evaporation to dryness and
dissolution of the residue. Measurements were taken using
a Perkin-Elmer DRC II ICP-MS.
Major compositions of minerals were measured on a
JXA-8900RL JEOL Superprobe equipped with wave-
length-dispersive spectrometers (WDS) and an energy
dispersive spectrometer (EDS). Operating conditions were
15.0 kV accelerating voltage, 12 nA beam current, and
5–10 lm electron beam diameter. Standards included were
as follows: wollastonite for SiO2 and CaO, synthetic TiO2
crystal for TiO2, synthetic Al2O3 crystal for Al2O3, syn-
thetic Cr2O3 crystal for Cr2O3, hematite for FeO, rhodonite
for MnO, synthetic MgO crystal for MgO, synthetic NiO
crystal for NiO, barite for BaO, albite for Na2O, and
sanidine for K2O.
Fig. 3 Semi-quantitative element mapping demonstrating the occur-
rence of amphibole (Am), chlorite (Chl), Cpx-I and Cpx-II, and
orthopyroxene (Opx) in the transition zone (CCSD-2-Tr). Cpx-I and
Cpx-II in the upper right map refer to primary and secondary
clinopyroxene, as defined in the text. Grt, garnet; Oli, olivine
Fig. 4 Compositional zoning profile of a Cpx-II grain from the
transition zone between the garnet peridotite and eclogite (sample
CCSD-2-Tr), showing increase in metamorphic temperatures (T/100)
and pressures (due to rimward Na2O increase)
Contrib Mineral Petrol (2011) 162:797–819 803
123
Trace element compositions of minerals were acquired
by in situ laser ablation inductively coupled plasma mass
spectrometry (LA-ICP-MS) both at the University of
Science and Technology of China and at the University of
Go¨ttingen. Both systems employ a 193-nm ArF Excimer
laser coupled to a Perkin-Elmer DRC II ICP-MS. All
samples have been prepared as polished thin sections
(ca. 2 mm thick). Instrumental background levels were
established by a gas blank, that is, analysis of the He–Ar
mixed gases with the laser off for 15–20 s. Laser ablation
was carried out in a He atmosphere mixed with Ar carrier
gas before the plasma torch. 29Si was used as an internal
standard; NBS 610 was used as a calibration standard.
Reproducibility and accuracy of trace element concentra-
tions were assessed to be better than 10% as checked with
both systems. Ablation spots were around 120 lm, with the
laser energy of 200 mJ. The ablated aerosol was carried to
the ICP source with He gas. Care was taken to perform
analyses on crack-free areas of the minerals. The combi-
nation of the high-quality optical imaging capabilities of
the laser-ablation system and the time-resolved analysis of
the ablation signals ensures that mineral inclusions were
excluded during measurements.
Li and B concentrations
Li and B concentrations in minerals were analyzed by in
situ LA-ICP-MS in which only Li and B as well as a few
inner-standard and monitoring elements were included in
order to increase the sensitivity during measurements. Li
and B concentrations were quantified using the SiO2 con-
tents of the corresponding electronic probe analysis. The
detection limit for Li is about 0.01 lg/g, sufficiently low to
investigate Li in the analyzed minerals.
The detection limit of B is not better than 1 lg/g,
because of possible surface contamination of the sam-
ples—although being carefully cleaned before measure-
ment—and the possibility of B liberation from the sample
tubing connecting the laser ablation cell with the ICP-torch
(e.g., Lee et al. 2008), Thus, concentrations of B \ 1 lg/g
in the analyzed minerals are estimates rather than true B
contents.
Lithium isotopes
Lithium concentrations and isotope ratios were determined
using a ThermoFinnigan MC-ICP MS Neptune at the
Research Centre for Earth Sciences in Potsdam (GFZ).
About 30–100 mg of powdered sample was dissolved in
HNO3/HF (1:5) and HClO4. Great care was taken to reach
complete dissolution, as any residue or precipitate is likely to
cause Li isotopic fractionation. Li was separated and purified
using a two-step ion-exchange chromatography. The first
separation step (HNO3 in methanol on AG 50 W-8 9
200–400 mesh) followed the setup of Wunder et al. (2006).
To remove Na completely, the eluate has been dried again
and redissolved in 0.5 ml of 0.5 N HCl in 80% methanol for
the second separation (see Jeffcoate et al. 2004). The sample
was ‘‘washed’’ with 2 ml of 0.5 N HCl in 80% methanol.
Lithium was evaluated from the column with 1 ml of 0.5 N
HCl in 80% methanol and with 3.5 ml of 1 N HCl. Complete
Li elution was checked by adding and collecting additional
0.5 ml of 1 N HCl. Wash solutions before and after Li elu-
tion were analyzed for their Li content. Li loss during the
chemical procedure was less than 1%. Each series of Li
separation included a NIST 8545 standard solution and ref-
erence material JR-2 or JG-2 to check the analytical proce-
dure. Analytical procedure and operating conditions closely
correspond to those of Wunder et al. (2006). They showed
that the Li blank can reach up to 130 pg but has no influence
on the isotopic composition. d7Li values are given relative to
Li reference material NIST RM 8545 (L-SVEC). The ana-
lytical precision and accuracy of d7Li were estimated by
repeated analysis of reference materials to be 0.19 ± 0.1%
(1sd) for JG-2 and 4.07 ± 0.34% (1sd) for JB-3, which are in
the range of published values (Wunder et al. 2006 and ref-
erences cited there in).
Boron isotopes
Boron concentration was determined on an aliquot of the
separated sample using a Varian Vista MPX simultaneous
ICP-OES, whereas B isotopic compositions were deter-
mined as Cs2BO2
? on a Finnigan MAT 262 (Nakano
and Nakamura 1998) at the GFZ Potsdam. About
500–1,000 mg of powered sample was fused with K2CO3
as described in Kasemann et al. (2001). Separation of
boron followed the three-step ion-exchange chromato-
graphic procedure developed by Tonarini et al. (1997). To
monitor the analytical procedure, IAEA reference material
B5, standard boric acid NBS 951, and a procedure blank
were included in each series. All samples with B contents
below 0.8 lg/g were analyzed in duplicate or triplicate
separate dissolutions. d11B values of the samples were
calculated relative to that of the reference material NBS
951 measured on the same turret. Data given in Table 4
have been corrected to NBS 951 that passed the entire
chemical separation procedure having a d11B value of
-0.05 ± 0.23 % (1sd) that reproduces the untreated
standard value. The precision and accuracy were estimated
based on the analysis of eight independent dissolutions
for IAEA standard B5 (basalt) that gave a d11B value of
-4.39 ± 0.53% (1sd), which is well in the range of pub-
lished data (Gonfiantini et al. 2003).
804 Contrib Mineral Petrol (2011) 162:797–819
123
Oxygen isotopes
Oxygen isotopes for minerals were analyzed by ‘‘in situ’’
UV-laser fluorination, whereas those for bulk rocks were
measured with a CO2-laser system in Go¨ttingen. Analyti-
cal procedures for UV-laser fluorination have been
described in detail in the study by Xiao et al. (2000) and
by Wiechert et al. (2002). CO2-laser measurements for
bulk rock were taken on 1–1.5 mg rock powder, which
was loaded in a nickel holder with eighteen 2-mm deep
holes. After evacuation and heating to *80C overnight,
the laser was operated at an energy level suitable to
transfer the powder lump into a ball, which avoids sample
loss during analysis. Afterward, the sample chamber was
loaded with 15–25 mbar F2, and the sample was heated
with the CO2-laser. Oxygen was purified in a way similar
to the UV laser method and directly transferred to the
Finnigan Delta plus for measurement of 18O/16O ratios.
Every 10 analyses were bracketed by an internal standard
(MORB glass), which has also been run at the beginning
and at the end of each measurement day. The routine
analytical error based on replicate analysis and standards
is better than 0.2%.
Sr, Nd, and Pb isotopes
Samples were dissolved in concentrated HF at 160C.
Digested samples were dried and taken up in 6 N HCl. Sr
and Nd were separated and purified as described in the
study by Romer et al. (2005). Sr isotopic composition was
determined on a VG54 Sector multi-collector mass spec-
trometer at GFZ Potsdam. Nd isotopic composition was
obtained on a Finnigan MAT 262 multi-collector mass-
spectrometer at the same laboratory. 87Sr/86Sr and
143Nd/144Nd were normalized to 86Sr/88Sr = 0.1194 and
146Nd/144Nd = 0.7219, respectively. Analytical uncertain-
ties are given at the 2 rm level. Pb was separated using a
HBr–HCl ion-exchange procedure (cf. Romer et al. 2005)
and analyzed on a Finnigan MAT 262 multi-collector
mass-spectrometer. Pb isotope data were corrected for
mass discrimination with 0.1% A.M.U. Reproducibility at
the 2r level is better than 0.1%.
Results
Major and trace element compositions of whole rocks are
listed in Table 3. Selected mineral microprobe and
LA-ICP-MS data are given in Supplementary Table 1 . Li
and B elemental and isotopic data are documented in
Table 4 and Fig. 6. Oxygen and Sr-Nd-Pb isotopic com-
positions are documented in Tables 5 and 6, respectively.
Trace elements and Li, B, O isotopes
Sample set 1
Trace element concentrations of the vein and the host
eclogite are much higher than those of the primitive mantle,
except high-field-strength elements (HFSE), which are
relatively depleted (Fig. 5a). Comparison of element con-
centrations in the vein and the host eclogite demonstrates
that CCSD-1-Hv is enriched in components released during
prograde dehydration. Thus, incompatible elements Rb, Sr,
and Ba, and LREE contents are enriched in CCSD-1-Hv
relative to CCSD-1-Ec, indicating that the HP vein is a sink
for fluid-mobile elements. By contrast, due to their relative
immobility, Nb, Ta, Ti, and HREE are depleted in the vein.
In addition, the two samples display very different REE
patterns: CCSD-1-Hv has much higher light REE contents
but lower heavy REE concentrations than CCSD-1-Ec
(Fig. 5a). Major element characteristic and trace element
partitioning between the host eclogite and the vein seem to
be the result of vein formation during prograde dehydration.
Of special interest are Li and B concentrations and
isotope compositions (Table 4). In eclogite, omphacite is
the phase with the highest Li concentration ranging from
39.26 to 41.60 lg/g, whereas the B content (0.66 lg/g) is
low. Phengite is the main carrier of B in the eclogite,
having B concentration from 5.17 to 5.42 lg/g. Garnet is
low in both Li and B concentrations, with 0.11 lg/g Li and
0.43 lg/g B. In the HP vein (CCSD-1-Hv), phengite has
slightly higher B and Li compared with the host eclogite,
with concentrations ranging from 5.47 to 6.13 and 0.64 to
0.73 lg/g, respectively. Eclogite has a d7Li of -1.6 %
(5.7 lg/g Li), whereas the HP vein (CCSD-1-Hv) has a
d7Li of 1.2% (2.2 lg/g Li) (Table 4; Fig. 6). Boron does
not show much difference between the two samples neither
in element concentrations nor in isotope compositions
(d11B values are -4.9 and -5.3%, respectively).
Bulk oxygen isotopic compositions (3.1%) in the vein
are higher than in the host eclogite (1.3%), resulting from
the higher modal abundance of quartz in the vein (Table 5).
Homogeneous d18O values around 2.8–3.0 % have been
measured in phengites from both the HP vein and the host
eclogite, suggesting that individual phengite grains grew in
equilibrium with the same fluid. Coarse-grained quartz
from the HP vein shows d18O values of 4.8–5.1. Estimated
temperature of vein formation, using quartz, and phengite
oxygen isotope thermometry (Zheng 1993) is around
750C.
Sample set 2
Garnet peridotite with a depleted heavy rare earth-element
(REE) pattern (normalized relative to the PM of Sun and
Contrib Mineral Petrol (2011) 162:797–819 805
123
Table 3 Whole-rock major and trace element data
Sample CCSD-1-Ec CCSD-1-Hv CCSD-2-Gp CCSD-2-Tr CCSD-2-Ec CCSD-3-Ec CCSD-3-Am CCSD-4-Ec CCSD-4-Av
SiO2 45.4 61.8 40.8 45.5 45.1 41.8 50.9 46.4 53.1
TiO2 2.19 0.60 0.94 0.49 1.66 5.34 0.91 2.47 2.34
Al2O3 18.9 14.8 5.9 11.8 13.3 13.3 15.5 14.5 10.8
Fe2O3 (Fetot) 12.8 4.4 14.2 13.3 14.7 18.0 9.9 19.1 10.1
MnO 0.25 0.07 0.19 0.27 0.28 0.31 0.15 0.35 0.08
MgO 6.71 4.63 29.6 15.6 13.2 5.55 6.50 4.62 5.80
CaO 8.79 5.87 5.20 9.77 9.76 11.9 8.10 8.76 10.5
Na2O 2.59 2.83 0.85 1.61 1.60 2.05 4.31 3.09 5.80
K2O 1.65 3.38 0.04 0.03 0.01 0.02 2.90 0.05 0.03
P2O5 0.07 0.41 0.02 0.03 0.02 2.92 0.24 0.48 0.45
Sum (wt%) 99.4 98.7 97.7 98.4 99.7 101.2 99.4 99.8 99.0
Li 5.70 2.20 7.50 12.0 3.60 6.87 4.70 18.00 7.70
Be 0.21 0.29 0.15 0.37 0.34 nd nd 1.25 0.58
B 1.15 1.77 4.51 1.51 0.91 0.50 1.30 0.40 0.40
Sc 24.7 7.89 9.94 21.4 39.8 14.2 18.3 27.4 43.3
V 146 169 140 254 314 131 205 439 303
Cr 29.0 25.5 1130 968 110 135 7.44 12.2 13.5
Co 17.4 9.27 130 96.4 37.9 26.8 25.4 21.2 28.4
Ni 62.6 37.1 997 873 62.3 64.5 5.16 68.5 74.4
Cu 4.59 1.99 55.7 119 6.39 16.3 17.2 14.0 12.0
Zn 64.1 48.1 86.1 30 32.7 90.3 85.4 111 96.8
Rb 28.6 61.4 0.58 0.43 0.12 0.41 31.3 0.76 1.07
Sr 61.6 189 129 291 279 90.6 412 202 162
Y 9.85 3.27 5.31 15.0 24.2 15.9 12.4 10.5 51.7
Zr 3.65 18.1 13.3 27.3 33.0 5.21 5.96 12.9 43.4
Nb 0.86 0.32 0.46 0.69 1.96 2.96 2.79 5.65 6.50
Mo 0.13 0.15 0.08 0.18 0.13 0.11 0.25 0.51 0.47
Cd 0.05 0.04 0.07 0.14 0.16 0.07 0.10 0.06 0.12
Sn 0.05 0.07 nd 0.33 0.83 0.67 0.37 0.09 0.07
Sb 0.04 nd 0.01 0.01 0.03 0.00 0.02 nd nd
Cs 0.64 1.41 0.10 0.14 0.05 0.43 0.00 0.03 0.02
Ba 241 662 12.2 50.3 44.2 7.88 354 35.6 21.3
La 0.38 2.39 1.87 4.11 3.98 1.93 18.0 16.7 21.8
Ce 0.98 7.05 5.69 12.7 12.3 4.76 42.1 42.2 52.8
Pr 0.16 1.04 0.89 2.01 1.97 0.67 6.25 5.74 6.90
Nd 1.04 5.75 4.32 9.75 9.8 3.51 31.2 26 30.7
Sm 0.66 1.58 1.10 2.48 2.58 1.09 7.65 6.06 7.17
Eu 0.58 0.75 0.38 0.85 0.93 0.42 2.42 1.67 2.28
Gd 1.25 1.19 1.22 2.86 3.3 1.82 7.61 4.31 7.01
Tb 0.29 0.15 0.18 0.43 0.56 0.39 0.97 0.57 1.40
Dy 1.85 0.70 1.08 2.8 4.13 3.16 5.30 2.51 9.29
Ho 0.40 0.13 0.19 0.54 0.88 0.74 0.93 0.42 1.98
Er 1.12 0.37 0.48 1.46 2.63 2.20 2.27 1.02 5.74
Tm 0.16 0.05 0.06 0.19 0.38 0.32 0.27 0.14 0.84
Yb 0.99 0.34 0.37 1.17 2.49 2.04 1.44 0.84 5.34
Lu 0.15 0.06 0.05 0.16 0.37 0.31 0.20 0.13 0.82
Hf 0.08 0.30 0.42 0.94 1.07 0.18 0.20 0.23 0.66
Ta 0.14 0.04 0.02 0.09 0.26 0.29 0.13 0.37 0.50
806 Contrib Mineral Petrol (2011) 162:797–819
123
McDonough 1989) has lower REE contents than both
eclogite and the transition zone (Fig. 5b). La/Yb ratios for
the garnet peridotite, transition zone, and eclogite are 5.1,
3.5, and 1.6, respectively. Garnet peridotite displays a
slight LREE enrichment relative to the depleted mantle
with a 2–3 times decrease in HREE (Fig. 5b). Eclogite
shows PM-normalized REE pattern comparable with
N-MORB. Such a pattern is in contrast to that of other
eclogites in ultramafic rocks from the Sulu area (Zhang
et al. 2009a). The transition zone (CCSD-2-Tr) shows
trace element patterns similar to the eclogite or those lying
between eclogite and garnet peridotite, except that it is
highest in Li but lowest in Ti of the three samples. An Eu
anomaly has been observed neither for the garnet perido-
tite nor for the eclogite. It is also noted that the garnet
peridotite and the transition zone have much higher Ni
contents (1,115 and 931 lg/g, respectively) than the
eclogite (66 lg/g; Table 3). Moreover, the garnet perido-
tite has a Nb/Ta ratio of 17.8, whereas the eclogite and the
transition zone have much lower ratios of 7.2 and 7.1,
respectively. In summary, major and trace elements sug-
gest distinct protolith sources for the peridotite and
eclogite.
Olivine in garnet peridotite contains 11.35–12.68 lg/g
Li but shows a large variation in B ranging from 3.70 to
18.9 lg/g (Table 4). Textural observations and electron
microprobe data indicate that olivine and serpentine in the
garnet peridotite are often intergrown. After careful
LA-ICP-MS spot analysis of pure olivine, we consider the
concentration of 3.70 lg/g as reflecting the B concentration
of the olivine; the higher concentrations are due to ser-
pentine intergrowth. The rarely found clinopyroxene in the
garnet peridotite has 15.30 lg/g Li and 4.12 lg/g B. What
is remarkable is that both olivine and clinopyroxene from
CCSD-2-Gp have much higher Li and B concentrations
than normally found in minerals of fresh orogenic perido-
tites (1.10–2.73 lg/g,Li and 0.04–0.28 lg/g B for olivine;
1.18–3.90 lg/g Li and 0.05–1.21 lg/g B for Cpx; Ottolini
et al. 2004).
Garnet peridotite (CCSD-2-Gp) shows very different
d11B values and B concentrations compared with eclogite
(CCSD-2-Ec). CCSD-2-Gp is enriched in B (4.5 lg/g) and
11B (d11B up to ?11.7 %) relative to CCSD-2-Ec (0.9 lg/g
and -0.1%, respectively), whereas the transition zone
(CCSD-2-Tr) has an intermediate composition, with
B = 1.5 lg/g and d11B = ? 5.8% (Table 4, Fig. 6).
Minor serpentine is intergrown with olivine in the garnet
peridotite (Fig. 2b). Thus, the measured B-content of
olivine in the garnet peridotite may be biassed by the
presence of serpentine. To distinguish the B content of
olivine from that of the serpentine, we checked after the
LA-ICP-MS measurements, the analyzed spots by electron
microprobe analyses. Contents of *3.70 lg/g B were
obtained for pure olivine, whereas up to 18.9 lg/g B was
found when *90% serpentine was involved. Considering
the high modal abundance of olivine and the much lower
one of the serpentine, we suggest that olivine is the major B
carrier phase, which accounts for more than 60% of the
total B inventory of the garnet peridotite. Therefore, the
systematic variation in B content and d11B probably
reflects fluid addition of B from the eclogite to the contact
garnet peridotite. In contrast to the large variation of B
isotopes, Li isotope variations in the three samples are
minor, being around 2% slightly lower than those for fresh
MORB (Chan et al. 1992; Elliott et al. 2002; Tomascak
et al. 1999).
Minerals from the garnet peridotite (CCSD-2-Gp) have
d18O ranging from 2.9 to 3.6 for olivine, 3.6–4.0 for
clinopyroxene, and 3.6–3.9 for garnet (Table 5), signifi-
cantly lower than those from unaltered mantle peridotite
(e.g., Chazot et al. 1997). Garnets and clinopyroxene from
the transition zone show d18O values of 3.1–3.3 and
3.3–3.6, whereas those from the eclogite have values of
2.7–3.1 and 3.2–3.4, respectively.
Sample sets 3 and 4
Within sample set 3, retrograde amphibolite (CCSD-3-Am)
has much higher fluid-mobile element concentrations (Rb,
Ba, Sr, 4–80 times higher) than eclogite (CCSD-3-Ec;
Table 3 and Fig. 5c) and displays much higher light REE
but similar heavy REE compared with the eclogite, despite
the much higher garnet abundance in the latter (Fig. 5c). A
characteristic feature of sample set 4 is that PM-normalized
Table 3 continued
Sample CCSD-1-Ec CCSD-1-Hv CCSD-2-Gp CCSD-2-Tr CCSD-2-Ec CCSD-3-Ec CCSD-3-Am CCSD-4-Ec CCSD-4-Av
W 0.12 0.13 nd 0.08 0.15 0.03 0.07 0.17 0.22
Tl 0.10 0.24 0.01 nd nd 0.23 nd nd nd
Pb 1.44 3.08 1.58 3.75 3.52 9.60 3.88 2.82 2.38
Th 0.07 0.35 0.03 0.06 0.16 0.28 1.23 1.52 2.43
U 0.03 0.06 0.01 0.01 0.02 0.12 0.45 0.37 0.49
nd not determined
Contrib Mineral Petrol (2011) 162:797–819 807
123
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808 Contrib Mineral Petrol (2011) 162:797–819
123
trace element patterns (Fig. 5d) are generally very similar
for eclogite and amphibolite, excluding HREE, which is
due to the higher garnet abundance in the host eclogite.
Sample CCSD-3-Ec (eclogite) has low B concentration
(0.5 lg/g) and a d11B value of -9.3%, whereas the ret-
rograde amphibolite (CCSD-3-Am) has higher B content
(1.3 lg/g) and a heavier d11B value (-1.1%) (Table 4;
Fig. 6). This suggests that B was added by fluids to the
amphibolite, and that, if the precursor of the amphibolite
had a similar B isotopic composition as the eclogite, the
fluid must have had strongly positive d11B values. Signif-
icant elemental and isotopic variations are also observed
for Li, with d7Li of -6.5% (14 lg/g) and of ?0.9%
(9.7 lg/g) in CCSD-3-Ec and CCSD-3-Am, respectively
(Table 4; Fig. 6).
For sample set 4, B concentration and isotopic compo-
sition are very similar, indicating limited B addition during
the formation of the retrograde amphibolite (CCSD-4-Av).
However, Li concentration and isotopic composition vary
between CCSD-4-Ec and CCSD-4-Av. The former has Li
Table 5 Oxygen isotopic compositions of bulk and major mineral phases in UHP rocks from the CCSD, Sulu, China
Sample Whole rocka Quartz Garnet Clinopyroxene Phengite Olivine
CCSD-1-Ec 1.26 (1.6b) 1.6–1.9 (4) 1.9–2.2 2.8–3.0 (3)
CCSD-1-Hv 3.07 (3.0b) 4.8–5.1 (5)c 1.9 (1) 2.4 (1) 2.8–3.2 (7)
CCSD-2-Gp 3.89 3.6–3.9 (2) 3.6–4.0 (3) 2.9–3.6(5)
CCSD-2-Tr 3.41 3.1–3.3 (3) 3.3–3.6 (6)
CCSD-2-Ec 3.08 2.7–3.1 (3) 3.2–3.4 (3)
CCSD-3-Ec 1.21
CCSD-3-Am 2.52
CCSD-4-Ec 1.68 1.7–1.8 (2) 2.0–2.1 (2)
CCSD-4;Av 1.79 4.2–4.4 (4) 1.7–1.9 (3)
a Values reported in the conventional d18O notation relative to the Standard Mean Ocean Water (SMOW)
b Value estimated from mineral oxygen isotopic compositions and modal abundance
c Number in bracket indicates the times of analyses
Table 6 Sr-Nd-Pb isotope data of investigated samples from CCSD, Sulu, Chinaa
Sample 87Sr/86Sr (±2 rm)
b 143Nd/144Nd (±2 rm)
c eNd
206Pb/204Pbd 207Pb/204Pb 208Pb/204Pb
Measured (T) Measured (T) Measured (T) Measured (T) Measured (T)
CCSD-1-Ec 0.708847 ± 7 0.70441 0.512794 ± 9 -2.7 17.402 17.35 15.492 15.49 37.906 37.87
CCSD-1-Hv 0.707813 ± 10 0.70471 0.512499 ± 5 -1.8 17.329 17.28 15.433 15.43 37.740 37.65
CCSD-2-Gp 0.705165 ± 10 0.70512 0.512364 ± 7 -4.1 16.904 16.89 15.392 15.39 37.313 37.30
CCSD-2-Tr 0.705114 ± 11 0.70510 0.512355 ± 5 -4.2 16.826 16.82 15.370 15.37 37.255 37.24
CCSD-2-Ec 0.705078 ± 7 0.70507 0.512417 ± 8 -3.3 16.804 16.79 15.371 15.37 37.268 37.26
CCSD-3-Ec 0.706328 ± 10 0.70632 0.512443 ± 5 -2.3 17.330 17.22 15.435 15.43 37.828 37.73
CCSD-3-Am 0.711226 ± 12 0.70793 0.512477 ± 5 -2.9 17.305 17.23 15.440 15.44 37.715 37.66
CCSD-4-Ece 0.706321 ± 7 0.70626 0.512350 ± 5 -3.9 17.809 17.32 15.486 15.46 38.437 37.65
0.706358 ± 7 0.70630 0.512356 ± 5 -3.8
CCSD-4-Ave 0.706292 ± 10 0.70626 0.512366 ± 5 -3.6 17.592 17.33 15.446 15.43 37.925 37.51
0.706295 ± 7 0.70626 0.512360 ± 5 -3.7
a The isotopic compositions of Sr, Nd, and Pb are from the same aliquots. The initial isotopic compositions were determined for T = 240 Ma,
which represents the time of the collision between the North China and Yangtze plates (Li et al. 1993). For analytical details, see Romer et al.
(2005)
b Analyzed using dynamic multicollection using a VG Sector 54-30 TIMS. Values are normalized to 86Sr/88Sr = 0.1194
c Analyzed using dynamic multicollection using a MAT 262 TIMS. Values normalized to 146Nd/144Nd = 0.7219
d Measured ratios corrected for mass fractionation with 0.1%/a.m.u. as determined from the repeated measurement of NBS 981 Pb reference
material. Reproducibility is better than 0.1%
e Sr and Nd isotopic compositions of samples CCSD-4-Ec and CCSD-4-Av were determined on separate dissolutions of aliquots from the same
powder
Contrib Mineral Petrol (2011) 162:797–819 809
123
content of 18 lg/g and a d7Li value of -6.9 %, whereas
the latter has 7.7 lg/g Li and a d7Li value of -0.4 %,
suggesting strong Li isotope fractionation between eclogite
and amphibolite (Table 4; Fig. 6).
Garnets from both CCSD-4-Ec and CCSD-4-Av have
very similar d18O values of 1.7–1.9, whereas quartz in
CCSD-4-Av has d18O values of 4.2–4.4 (Table 5). Coex-
isting garnet–quartz pairs (Zheng 1993) in the vein yield
0.01
0.1
1
10
100
CCSD-1-Hv CCSD-1-Ec
0.01
0.1
1
10
100
CCSD-2-Gp CCSD-2-Tr
CCSD-2-Ec
0.01
0.1
1
10
100
CCSD-3-Am
CCSD-3-Ec
0.01
0.1
1
10
100
CCSD-4-Ec
CCSD-4-Av
0.01
0.1
1
10
100
0.01
0.1
1
10
100
CCSD-2-Tr
0.01
0.1
1
10
100
0.01
0.1
1
10
100
Fig. 5 Primitive mantle-normalized bulk rock trace element patterns for selected samples from the CCSD. Normalized values for primitive
mantle are from Sun and McDonough (1989)
Fig. 6 d7Li versus Li (a), d11B
versus B (b), d7Li versus d11B
(c), and B/Nb versus d7Li
(d) plots for investigated UHP
metamorphic rocks from the
CCSD
810 Contrib Mineral Petrol (2011) 162:797–819
123
equilibrium temperatures of about 850C, indicating oxy-
gen isotope equilibrium under high-temperature conditions.
Whole-rock Sr-Nd-Pb isotopes
The Sr-Nd-Pb isotopic data are summarized in Table 6. Sr
and Nd isotopic compositions show large 87Sr/86Sr
(T = 240Ma) variations ranging from 0.70441 to 0.70793 and
low eNd values from -2.3 to -4.2 at 240 Ma, representing
the age before subduction (Li et al. 2000). In a 87Sr/86Sr(240
Ma) versus eNd plot (not shown), all investigated samples
coincide with eclogites from the Dabie–Sulu area, which
are characterized by relatively enriched Sr and Nd isotopic
compositions (John 1998; Wawrzenitz et al. 2006). The
sample CCSD-1-Ec shows a slight difference in measured
but similar initial (calculated for 240 Ma) 87Sr/86Sr ratios
and eNd values to the HP vein (CCSD-1-Hv), indicating that
vein-forming fluids were probably derived from the
eclogite (e.g., Becker et al. 1999).
The three samples from the coexisting garnet peridotite
and eclogite have similar initial Sr and Pb isotopic com-
positions, but slightly different initial Nd isotopic compo-
sitions (Table 6). A high 87Sr/86Sr ratio and a negative
initial eNd value in the garnet peridotite (CCSD-2-Gp) may
reflect input from subducted eclogites (CCSD-2-Ec).
As typical for eclogites of the Dabie–Sulu area, all
analyzed samples have unusually unradiogenic Pb isotopic
compositions due to a marked reduction in their Th/Pb
and U/Pb ratios during a Palaeoproterozoic or Archean
high-grade metamorphic event (Li et al. 2003; Wawrzenitz
et al. 2006). In 207Pb/204Pb(T) versus
206Pb/204Pb(T) and
208Pb/204Pb(T) versus
206Pb/204Pb(T) (T = 240 Ma) diagrams
(not shown), coexisting garnet peridotite and eclogite fall
in the field of post-collisional mafic-ultramafic intrusive
(PCMI) rocks (Li et al. 2009). The similarity in unradio-
genic Pb isotopic compositions of eclogite and peridotite
(Table 6) suggests that the Pb budget of the garnet peri-
dotite is dominated by eclogite-derived Pb,
Retrogressed amphibolite and eclogite, CCSD-4-Ec and
CCSD-4-Av, have very similar initial Sr and Nd isotopic
compositions. On the other hand, significant differences in
Sr isotopes are observed between CCSD-3-Ec and CCSD-
3-Am (Table 6), which suggest Sr input from the retro-
grade fluid. The very similar initial Nd and Pb isotopic
compositions of CCSD-3-Ec and CCSD-3-Am suggest that
the rocks had the same precursor lithology.
Discussion
The 4 sets of samples characterize 3 different temperature–
pressure conditions in a continental subduction-related
environment. We will discuss the 3 metamorphic settings:
(i) element transport during prograde dehydration condi-
tions, (ii) material exchange among mantle- and slab-
derived rocks and, (iii) fluid/rock interactions during
retrograde metamorphism. Of special importance for the
discussion is the role of Li and B during the various
metamorphic stages.
Element transport during prograde dehydration
Before we discuss the evidence for element transport under
prograde dehydration, we have to consider whether the
vein represents segregated partial melt. Peak metamorphic
conditions of the investigated HP vein up to *800C and
up to at least 3.5 GPa, close to the water-saturated solidus
of eclogite (Schmidt et al. 2004b), should favor melting.
However, the preservation of quartz and phengite in the
residual eclogite suggests that the HP vein cannot be a melt
phase from the host eclogite. In particular, the sharp con-
tact between the vein and the eclogite speaks against partial
melting. Furthermore, the SiO2 content in the vein
(61 wt%) is too low for a partial melt of a mafic rock at
eclogite facies conditions (Rapp and Watson 1995). In
addition, the enrichment of phengite, coarse-grained
quartz, and amphibole in the vein suggests that the vein
most likely was deposited from silicate-rich fluids (Kessel
et al. 2005; Schmidt et al. 2004a).
Phengite occurs in both the HP vein and the host
eclogite. Its major element (such as Si and K) and trace
element composition (Rb and REE) is comparable in both
phases. Calculations based on available thermobarometers
yield similar peak metamorphic pressures of 3.4–3.5 GPa
and temperatures of *800C for both the vein and the host
eclogite, which indicates that phengites have been subject
to the same P–T evolution. However, phengite from the HP
vein has much higher Sr and Ba and slightly higher Eu
concentrations relative to the host eclogite, indicating that
fluid-mobile trace elements have been transported from the
host eclogite to the HP vein.
As shown in Fig. 7, SiO2 content of phengite from
CCSD-1-Hv is negatively correlated with BaO content: the
phengite core has low Si and high Ba contents, whereas the
rim shows higher Si but lower Ba concentrations. Few
analyzed spots from the outermost phengite rim have both
low SiO2 and BaO contents, which may indicate that the
rims had originally higher Si contents being lowered during
retrogression. As mineral cores record earlier growth
histories than rims, the increase in Si from core to rim
indicates increasing pressures during phengite growth (e.g.,
Massone and Schreyer 1987). Since Ba is a typical fluid-
related element, the correlation suggests that the phengite
cores equilibrated with a Ba-rich fluid. Overall, these
observations suggest that phengites in the vein formed
under a typical subduction dehydration environment.
Contrib Mineral Petrol (2011) 162:797–819 811
123
Garnet and omphacite from the HP vein have compa-
rable middle to heavy REE contents but higher Rb, Sr, Ba,
and LREE concentrations relative to the host eclogite
(Supplementary Table 1) that may imply material transport
from the host eclogite into the vein. Whole-rock oxygen
isotope calculation based on mineral modal abundances
and measured mineral d18O values yields a bulk value of
3.0 % for the HP vein and 1.6 % for the host eclogite. This
difference is due to the different modal abundance of
quartz (d18O = *5.0). Therefore, trace element, mineral-
ogical and oxygen isotope data suggest that the fluid
responsible for the formation of the HP vein originates
from the surrounding eclogite. Comparing the volume of
the vein with that of the host eclogite, fluids forming the
vein must not necessarily be derived from the immediate
host rock, but instead, may originate from a larger portion
of surrounding dehydrated eclogites.
As Li and B are fluid-mobile elements (Marschall et al.
2006), they will be expelled during dehydration causing
isotope fractionation. Since the heavy isotopes 7Li and
11B preferentially enrich in the fluid phase (Benton et al.
2004; Benton et al. 2001), the deposited minerals in the
HP vein should be depleted in 7Li and 11B. However,
compared with the host eclogite, the HP vein has higher
d7Li but lower d11B values, reflecting possibly different
behavior of the two elements during vein formation.
Apparently, a major part of the fluid and some of it
contained Li and B that are mobile at relatively low P–T
conditions would migrate away from the vein-forming
fluid, lowering significantly the element concentrations
and probably also changing the isotopic compositions in
the vein minerals. The Li and B elemental and isotopic
features of the vein will thus not reflect the budget of the
dehydrated fluids but that of a residue after a series of
complex processes.
Material and fluid exchange between
mantle- and slab-derived rocks
Fluids released from subducting oceanic crust (Bebout
et al. 1993; Bebout and Barton 2002; King et al. 2003;
Kohler et al. 2009) or from continental crust (Marocchi
et al. 2007, 2009; Zhang et al. 2007) may transfer large
amounts of volatiles and trace elements to the overlying
mantle wedge. The metasomatic mineral assemblage at the
interface between peridotite and eclogite represents one
rare example in orogenic peridotite that has been affected
by UHP metamorphic conditions. The presence of amphi-
bole and chlorite in the transition zone between garnet
peridotite and eclogite is strong evidence for a metasomatic
process in which fluids have been added to garnet perido-
tite. Before we discuss this topic in more detail, we have to
address the question: whether garnet peridotite and eclogite
are rocks from different sources brought together during
slab subduction or alternatively products of differentiation,
i.e., cumulates of an intruded ultramafic body.
Primitive-mantle-like REE patterns (Fig. 5b) in the
garnet peridotite (CCSD-2-Gp) favor a mantle origin for the
rock. As mentioned, the investigated garnet peridotite has
been collected from the drill hole ZK 703 that is *80 m
away from the main hole of the CCSD that penetrates the
same ultramafic rock body (Xiao et al. 2006). Based on a
general petrochemical study on the recovered garnet peri-
dotite from the CCSD main hole, Yang et al. (2007) argued
that garnet peridotite and interlayered eclogite should
derive from different sources. Such a conclusion is sup-
ported by different formation ages of the two rock types:
peridotites have protolith ages of 346–461 Ma (Yang et al.
2007), whereas eclogites interlayered with the peridotite
have protolith ages[600 Ma (Zhang et al. 2009b). Further
support is obtained from mineral chemistries at the interface
between eclogite and garnet peridotite. As shown in Fig. 3,
in the transition zone, two compositionally different Cpx
grains (Cpx-I and Cpx-II) may occur, with Cpx-II as
overgrowths at the rim of Cpx-I or as individual grains
showing compositional core-rim zoning. Profiles through
Cpx-II grains show decreasing MgO and CaO contents
whereas jadeite component and Al2O3 increase from core to
rim of the grain (Fig. 4). Assuming chemical equilibrium
between the rims of coexisting Cpx-II and garnet, Cpx–Grt
thermometry of Aranovich and Pattison (1995) yields
temperatures of 960–1,020C at the rim and 780–800C at
the core for Cpx-II (Table 2; Fig. 4). Such zoning pat-
terns—also observed in orthopyroxene grains—may have
been obtained during prograde metamorphism. The zoned
Cpx-II and Opx grains indicate near-peak formation con-
ditions, suggesting that coexisting amphibole and sulfide
(FeS) in the transition zone (Fig. 2f) might have formed
under similar P–T conditions.
0
0.5
1
48 50 52 54
SiO2(wt%)
B
aO
(w
t%
)
Core
Mantle
Rim
Phengite
in CCSD-1-Hv
Fig. 7 SiO2 versus BaO profiles through coarse-grained phengite
from the high-pressure vein (CCSD-1-Hv), divided into core (blue),
mantle (pink), and rim (yellow) areas
812 Contrib Mineral Petrol (2011) 162:797–819
123
In summary, the sharp contact between the two rock
types (Fig. 2b), the presence of hydroxyl-bearing minerals
along the interface, a[10% bulk MgO difference between
peridotite and eclogite, and the different protolith ages
confirm that garnet peridotite and eclogite represent rocks
from different sources that have been brought together
during slab subduction. Numerical modeling by Gerya and
Sto¨ckhert (2006) has shown that rocks from different
lithospheric levels, following different P–T paths, can be
amalgamated through a ‘‘subduction channel.’’ The rim-
ward temperature and pressure increases in Opx and Cpx-II
in the transition zone suggest that the garnet peridotite was
brought together with the down-going eclogite before peak
metamorphism.
The growth of new minerals such as amphibole, chlorite,
and Cpx II strongly suggests formation by hydrous fluids
rich in silicate component. Comparison of mantle-nor-
malized trace element patterns of garnet and clinopyroxene
in garnet peridotite and in eclogite shows the characteristic
features of associated subduction zone fluids, which is also
documented by the overall enrichment of Rb, Ba, and Sr in
the garnet peridotite minerals (Supplementary Table 1).
Garnet peridotite (CCSD-2-Gp) has bulk Li and B contents
of 7.5 and 4.5 lg/g, respectively, which are considerably
higher than previously reported values for the depleted
mantle (typically \2 and \1 lg/g, respectively) (Palmer
and Swihart 1996; Paquin and Altherr 2002; Paquin et al.
2004; Sun and McDonough 1989). Studies on oceanic
peridotite showed that B is generally enriched, whereas Li
is depleted in serpentine minerals and serpentinite during
alteration (e.g., Niu 2004; Vils et al. 2008). As the garnet
peridotite is partially serpentinized, it might be argued that
the B enrichment results from the late serpentinization.
However, in situ measurement of B and Li concentrations
indicates that primary olivine and clinopyroxene have B
contents up to 3.7 and 4.1 lg/g, respectively, which are
much higher than those from depleted mantle (Ottolini
et al. 2004; Salters and Stracke 2004). The B enrichment,
thus, seems to be a primary feature of the rock. As shown
by the narrow range of Li concentrations in the analyzed
spots, the slight serpentinization does not affect signifi-
cantly the Li concentrations in the primary mineral phases.
Therefore, B and probably Li enrichment in the garnet
peridotite results from metasomatic fluids released during
the dehydration of the subducting slab (represented by the
eclogite).
In contrast to earlier suggestions that B cannot reach
depths greater than the volcanic roots of arcs (Ishikawa and
Nakamura 1994; Morris et al. 1990), the present study
indicates that a relatively large amount of crustal B and Li
can be recycled to mantle depths. This is in line with You
et al. (1993), who estimated that possibly less than 30% of
slab B input from subducted sediments and altered oceanic
crust is brought back to the surface in arcs. This is also
consistent with the modeling results by Marschall et al.
(2007b) who suggested that about 55% Li and *37% B in
K2O-bearing MORB may retain in the rocks up to eclogite-
facies metamorphism. Therefore, crustal-derived Li and B
may contribute significantly to the Li and B budget of the
mantle.
With respect to isotopes, Li shows a small isotopic
fractionation from 1.8 to 2.5 % relative to typical mantle
values (Chan et al. 1992; Elliott et al. 2002, 2004;
Tomascak 2004), whereas B displays significant isotopic
fractionation along the profile from eclogite to garnet
peridotite, with d11B values of -0.1 for eclogite to 5.8 for
the transition zone and to 11.7 for garnet peridotite
(Table 4; Fig. 6). It is unlikely that the B isotope frac-
tionation results from late-stage, low-temperature pro-
cesses, because olivine and clinopyroxene are the major
B-bearing phases (Table 4) and because the two rocks
behaved structurally coherent during exhumation. Any
fluid/rock interaction at low temperatures is likely to affect
both the Li and B isotope systems. The contrasting
behavior of Li and B under mantle conditions may be due
to that: (i) Li isotope fractionations are smaller than B
(James et al. 2003; Kisakurek et al. 2004; Rudnick et al.
2004; Zack et al. 2003); (ii) fluids released from UHP
eclogites may have much lower eclogite/fluid partition
coefficients for B than for Li (Kessel et al. 2005); and
(iii) the whole-rock budget of Li and B is controlled by dif-
ferent minerals, i.e., B is preferentially hosted in phengite
(Domanik et al. 1993; Marschall et al. 2006; Wunder et al.
2005; this study), whereas Li is predominantly hosted in
clinopyroxene and olivine (Marschall et al. 2006; Seitz et al.
2004; Woodland et al. 2002; Zack et al. 2002). Thus, the
small range of d7Li values may indicate an exchange at very
high temperatures.
Oxygen isotope data are in line with these arguments.
The small increase in d18O values from 3.1 for the eclogite
to 3.3 for the transition zone and to 3.9 for the garnet
peridotite indicates an addition of fluids from the eclogite
to the garnet peridotite, because the isotopic front of the
fluids from the former (with d18O of *3.1%) to the latter
(with d18O[3.9%) will become less steep with increasing
distance from the eclogite. Those fluids are released from
the ecogite is also documented by the occurrence of chlo-
rite and amphibole in the transition zone between the two
rocks.
Fluid/rock interaction during exhumation:
origin(s) of the retrograde fluids
With respect to the source of fluid(s) that caused the for-
mation of retrograde amphibolite in the Dabie–Sulu area,
numerous studies ascribed the origin of retrogression to
Contrib Mineral Petrol (2011) 162:797–819 813
123
deep infiltration of external fluids (Baker et al. 1997; Xiao
et al. 2000; Yui et al. 1997), while others argued that the
liberation of internal fluids from the rocks themselves
could be an important source (Chen et al. 2007; Li et al.
2004; Zheng 2009; Zheng et al. 2007). In the present case,
an external fluid source can explain the very different
fluid-mobile element budgets as well as oxygen isotopic
compositions in samples CCSD-3-Ec and CCSD-3-Am;
however, internal fluids released by the exsolution of
structural water from nominally anhydrous minerals after
peak metamorphism may provide sufficient amounts of
aqueous fluid for amphibolite-facies retrogression (e.g., Li
et al. 2004; Chen et al. 2007; Zheng et al. 2007; for review,
see Zheng 2009) in CCSD-4-Av. This is indicated by the
observation that CCSD-4-Av has a 40% higher modal
abundance of water-rich omphacite than CCSD-4-Ec (see
below). Thus, both external and internal fluids may con-
tribute to the retrogression of the UHP metamorphic rocks.
The investigated retrograde amphibolite samples
(CCSD-3-Am and CCSD-4-Av) may have different fluid
sources, as indicated by their different major components
and trace elements characteristics. The markedly higher
contents in Fe, Ti, and P in CCSD-3-Ec may be due to
cumulate ilmenite and apatite in the precursor. The 4–80
times higher fluid-mobile element concentrations of Rb,
Ba, and Sr in CCSD-3-Am relative to CCSD-3-Ec suggest
that a substantial volume of these elements must have been
added from external fluids (Fig. 5c), consistent with the
occurrence of fractures within the amphibolite (Fig. 1c).
The fact that no structural break is observed between
CCSD-3-Ec and CCSD-3-Am implies that the mineralog-
ical difference between the two samples largely reflects the
contrasting extent of retrograde overprint.
On the other hand, samples CCSD-4-Av (amphibolite-
vein) and CCSD-4-Ec (host eclogite) are similar in TiO2,
MgO, and P2O5 contents. Using bulk rock major element
data and measured compositions of omphacite and garnet,
an eclogite composed of 88% omphacite and 12% garnet is
consistent with sample CCSD-4-Av, whereas an eclogite
with 48% omphacite and 52% garnet may explain the
concentrations of major elements in CCSD-4-Ec (Fig. 8).
The presence of about 2% rutile in garnet and omphacite
from CCSD-4-Ec explains the similar TiO2 contents in
both samples. Furthermore, the two samples are also very
similar in most trace element concentrations (Table 3;
Fig. 5d), including the fluid-mobile elements, except the
HREE, which are much higher in the host eclogite
(Fig. 5d). Therefore, the formation of CCSD-4-AV results
from retrogression of omphacite to amphibole in an
omphacite-enriched layer, as omphacite may have much
more structural water than garnet (Chen et al. 2007) and
react easier with fluid relative to garnet.
This model again is consistent with oxygen isotope data.
For sample set CCSD-3, eclogite CCSD-3-Ec has a bulk
d18O value of 1.3, whereas the retrograde amphibolite
CCSD-3-Am has a d18O-value of 2.5. Considering the
mineral assemblage of the two rocks, in which 18O-enri-
ched minerals such as quartz are absent or minor, it seems
likely that the d18O variation between the amphibolite and
eclogite is the result of an external fluid/rock interaction
during the retrogression of CCSD-3-Am. On the other
hand, almost identical d18O values of both garnet and bulk
rock between CCSD-4-Ec and CCSD-4-Am, and equilib-
rium oxygen isotope fractionation between garnet and
amphibole in CCSD-4-Am (with mean d18O = 1.8 and 2.1,
respectively) indicate an internal fluid supply that caused
the formation of the amphibolite, as internal fluids related
to subduction processes usually do not change their d18O
[0.2% compared with their precursor rocks (Cartwright
and Barnicoat 1999).
Li and B behavior during retrograde metamorphism
of UHP metamorphic rocks
Major and trace elements, oxygen, and Sr-Nd-Pb isotopic
compositions indicate that sample sets 3 and 4 have formed
in different environments, i.e., external and internal fluid/
rock interaction systems, respectively. Therefore, Li and B
isotopes provide an opportunity to investigate high-tem-
perature Li and B isotopic fractionations under open- and
closed-system conditions during exhumation of UHP
metamorphic rocks.
Li and B isotopes are strongly fractionated between
CCSD-3-Am and CCSD-3-Ec (open system), with
Fig. 8 Comparison between
measured bulk major
components and those
calculated from microprobe data
of omphacite and garnet in
sample set 4. Modeling suggests
that CCSD-4-Av composes of
88% omphacite ? 12% garnet,
whereas CCSD-4-Ec contains
48% omphacite and 52% garnet
814 Contrib Mineral Petrol (2011) 162:797–819
123
d7Li = -6.5 and d11B = -9.3 in the eclogite, and sig-
nificantly heavier values (?0.9 and -1.1, respectively) in
the retrograded amphibolite, indicating an enrichment of
7Li and 11B in the amphibolite having reacted with an
external fluid. Moreover, higher Li but lower B concen-
trations in the eclogite compared with the amphibolite
(Table 4) indicate that, B has been added during fluid/rock
interaction, whereas Li is more controlled by mineral
characteristics rather than by the fluid alone, which is
reflected by the much higher Cpx abundance in the eclogite
than in the amphibolite.
In CCSD-4-Am and CCSD-4-Ec, representing an inter-
nally buffered fluid/rock system, the very similar B element
concentration and isotopic compositions of eclogite and
amphibolite indicate that B concentration and isotopic
composition will not change significantly when the fluid is
internally derived. On the other hand, Li concentration and
isotopic composition are significantly different between
eclogite and amphibolite (Table 4), suggesting significant
Li elemental and isotopic fractionation under closed con-
ditions. This again indicates that under high P–T conditions
(amphibolite formation: temperatures are around 800C),
mineralogical differences rather than fluids will affect the
Li budget and isotopic compositions of a rock.
Although at low temperatures (e.g., \300C) both Li
and B have similar small solid/fluid partition coefficients
and the same isotopic fractionation trends, at higher tem-
peratures, B behaves as a typical fluid-mobile element
(Leeman and Sisson 1996), whereas Li concentration and
isotopic composition are more reaction controlled than by
fluid phases. During high-grade metamorphism (i.e., high
T), Li and B thus may show different patterns of enrich-
ment and of isotopic fractionation.
Conclusions
Four sample sets with contrasting lithological boundaries
in UHP metamorphic rocks from the CCSD program pro-
vide information on element transfer and Li, B, and O as
well as Sr-Nd-Pb isotopic features induced by fluid mobi-
lization from subduction dehydration through peak meta-
morphism to exhumation rehydration processes in a
continental subduction zone. During the various stages of
UHP metamorphism, O and Sr-Nd-Pb isotopic composi-
tions keep in general unchanged, but there is significant Li
and/or B isotope fractionation between different rock types
that are in close contact. An enrichment of Rb, Ba, Th,
LREE, and other incompatible elements in a HP vein
representing dehydration during prograde metamorphism is
interpreted that vein formation is initiated by locally
derived fluids that have leached elements from the dehy-
drated host rock. These fluids are water rich, with low
contents of silicates at an early stage, but under UHP
conditions become phases that are intermediate in com-
position between water and silicate melt. Partial melting is
unlikely even under high water activity of vein-forming
conditions, as partial melting will remove any phengite in
the vein.
Due to lithological and structural heterogeneities of the
subducting slab, the temporal and spatial distribution of
fluids is heterogeneous within the slab. This is documented
in a sample set consisting of garnet peridotite, a transition
zone, and eclogite. The formation of chlorite and amphi-
bole and compositionally different Cpx-II grains in the
transition zone between eclogite and garnet peridotite
clearly reveals the addition of hydrous fluids enriched in
silicate component from the subducting slab to the mantle.
Investigations into retrogressive samples revealed that
both external and internal fluids are plausible candidates
that may have caused retrogression of UHP metamorphic
rocks in the Dabie–Sulu area, depending on open or closed
system conditions.
The overall large variations of d7Li from -6.9 to ?2.5%
and d11B from -9.3 to ?11.7 in the investigated samples
indicate significant Li and B isotopic fractionations during
different metamorphic stages of the UHP rocks. Regardless
of the origin of the d7Li and d11B variations observed in
four sample sets, two important points should be empha-
sized. First, B and Li may behave differently during UHP
metamorphism as their budget is probably controlled by
fluid, melt, temperature, and mineral assemblages as well as
diffusive processes, and most importantly by the previous
history of fluid and mineral reactions. This is evident from
the contrasting isotopic fractionation of Li and B in samples
representing prograde metamorphism and those represent-
ing near-peak conditions and the contrasting correlation
between element concentration and isotopic composition of
the two elements. Second, Li and B isotopic compositions
of retrograde amphibolite that may result from external
fluid/rock interactions are consistent with experimental data
on the fractionation behavior of Li and B among minerals
and HP fluids. However, in a closed system (e.g., sample set
4), mineral assemblages seem controlling Li isotopic
fractionations.
Acknowledgments Special thanks go to Dr. R. L Romer for his
support in the analytical work. He introduced the first author to the
radiogenic isotope (Sr, Nd, Pb) chemistry and performed the Sr-Nd-
Pb isotope analysis. Mrs. A. Meixner was of great help for the Li and
B isotope analysis and carried out the MC-ICPMS measurements.
Analytical assistance from Mrs. C. Schulz is highly appreciated. Prof.
ZQ Xu is thanked for her assistance during sample collection. Careful
and constructive reviews by Horst Marschall and an anonymous
reviewer and editorial comments by Tim Grove have substantially
improved the manuscript. This study was financially supported by
grants from the National Science Foundation of China (40773003,
40921002), the National Science Foundation of Germany (HO
Contrib Mineral Petrol (2011) 162:797–819 815
123
375/22) the Academy of Science of China (kzcx2-yw-131), and the
Major State Basic Research Development Program of China (2009
CB825002, 2003 CB716501).
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