Phase Equilibrium Modeling of MT–UHP Eclogite: a Case ...Yangkou Bay in the Sulu Belt of eastern...
Transcript of Phase Equilibrium Modeling of MT–UHP Eclogite: a Case ...Yangkou Bay in the Sulu Belt of eastern...
Phase Equilibrium Modeling of MT–UHP
Eclogite: a Case Study of Coesite Eclogite at
Yangkou Bay, Sulu Belt, Eastern China
Bin Xia1,2*, Michael Brown2,3, Lu Wang1,3, Song-Jie Wang1,2,3,4 and
Philip Piccoli2
1School of Earth Sciences, State Key Laboratory of Geological Processes and Mineral Resources, China
University of Geosciences, Wuhan 430074, China; 2Department of Geology, Laboratory for Crustal Petrology,
University of Maryland, College Park, MD 20742, USA; 3Center for Global Tectonics, China University of
Geosciences, Wuhan 430074, China; 4College of Earth Science and Engineering, Shandong University of Science
and Technology, Qingdao 266590, China
*Corresponding author. Present address: College of Earth Sciences, State Key Laboratory of Geological
Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China. E-mail:
Received December 3, 2016; Accepted May 28, 2018
ABSTRACT
In this study, we present an example of phase equilibrium modeling of medium-temperature–ultra-
high-pressure (MT–UHP) eclogites that includes consideration of the influence of ferric iron (O) and
H2O on the phase equilibria. As a case study, we focus on the intergranular coesite-bearing eclogites atYangkou in the Sulu Belt. Based on phase equilibrium modeling of four eclogites, we monitor changes
in phase relations during deep subduction and exhumation, and investigate fluid behavior during de-
compression. To determine the appropriate O and H2O contents to use in calculating P–T pseudosec-
tions for these eclogites, we use an iterative process in which calculated temperature/pressure (T/P)–O/
H2O phase diagrams are combined with constraints from petrological observations. P–T pseudosec-
tions were calculated for each of the four eclogites to constrain the P–T conditions. The highest P–T con-
ditions retrieved were P> 5�5 GPa at T>850�C, although variation in mineral compositions suggeststhat the maximum P–T conditions could have been higher. A P–T path was reconstructed based on
microstructural evidence, mineral compositions that constrain P–T conditions within phase assemblage
fields, average P calculations and mineral thermobarometry. During exhumation, the retrograde P–T
path passed through metamorphic conditions of P¼ 4�0–3�4 GPa at T¼ 850–800�C and P¼2�4–1�7 GPa
at T¼ 800–750�C, before reaching crustal levels at P¼ 1�3–0�9 GPa at T¼ 730–710�C. The prograde evo-
lution is suggested to have followed a high dT/dP path during the early stage of subduction, followedby a low dT/dP segment to the metamorphic peak. During exhumation, the eclogites at Yangkou be-
came domainal, made up of host-rock with low a(H2O) in which garnet and omphacite have partially re-
equilibrated and intergranular coesite has been preserved, cut by veins and veinlets where a(H2O) was
higher and new mineral assemblages have developed. In the veins, the new assemblage comprises
coarse phengite and quartz with symplectites of K-feldspar þ plagioclase þ biotite þ quartz around the
phengite. By contrast, the veinlets comprise symplectites of hornblendeþ plagioclase 6 quartz 6 clino-
pyroxene after omphacite; similar symplectites occur at the edges of the phengite–quartz veins againsthost eclogite. We interpret the coarse phengite and quartz, which previously could have been coesite,
to have formed by precipitation of solutes from fluid migrating under UHP conditions, whereas we in-
terpret the symplectites around the phengite to have formed by local melting and crystallization during
exhumation from HP eclogite- to HP amphibolite-facies conditions. The symplectites in the veinlets and
along the edges of the phengite–quartz veins are interpreted to have formed by reaction of local grain-
boundary fluid with the host under HP amphibolite-facies conditions.
VC The Author(s) 2018. Published by Oxford University Press. All rights reserved. For permissions, please e-mail: [email protected] 1253
J O U R N A L O F
P E T R O L O G Y
Journal of Petrology, 2018, Vol. 59, No. 7, 1253–1280
doi: 10.1093/petrology/egy060
Advance Access Publication Date: 11 June 2018
Original Article
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Key words: intergranular coesite; MT–UHP eclogite; phase equilibrium modeling; P–T path;Yangkou, Sulu Belt
INTRODUCTION
Ultrahigh-pressure (UHP) metamorphic rocks, particu-
larly eclogite and associated country rock gneisses in
orogenic belts, demonstrate that continental crust can
be subducted to and returned from mantle depths
(Chopin, 2003; Liou et al., 2004; Brown & Johnson,
2018). Pressure–temperature–time (P–T–t) paths tracing
the deep subduction and exhumation of these UHP
rocks provide insight into geological processes during
continental collision at convergent plate boundaries
and form the basis for geodynamic modeling of these
processes (Gerya & Stockhert, 2006; Warren et al.,
2008; Roda et al., 2012; Sizova et al., 2012). Thus, robust
quantification of P–T–t paths from natural samples of
UHP metamorphic rocks is important if our geodynamic
modeling is to provide deeper understanding of proc-
esses during continental collision at convergent plate
boundaries.
Although the P–T conditions of UHP metamorphism
can be qualitatively constrained by the presence of indi-
cative UHP minerals such as coesite and diamond, the
quantitative estimation of these conditions is under-
taken using conventional thermobarometry (e.g. Krogh
Ravna & Terry, 2004) and/or phase equilibrium model-
ing (e.g. Wei & Clarke, 2011; Wei et al., 2013). Forward
modeling involves the calculation of phase equilibria
for a given rock composition using an internally consist-
ent thermodynamic dataset and appropriate activity–
composition models for the phases of interest (Holland
& Powell, 1998; Powell et al., 1998), which may then be
related to the observed mineral assemblages, mineral
proportions and mineral compositions for that particu-
lar sample. In addition, we may calculate phase equili-
bria for a representative composition (e.g. mid-ocean
ridge basalt; MORB) to investigate how variables such
as H2O content and oxidation state affect these equili-
bria (Rebay et al., 2010).
During the last decade, phase equilibrium modeling
has become the preferred thermobarometric method in
many studies because it utilizes the maximum informa-
tion available from the sample being studied, and, in
many cases, allows the evolution of mineral assemb-
lages to be quantified to determine a robust P–T path
(Powell & Holland, 2008). This method has proven use-
ful in the study of UHP eclogites, in part because some
minerals, such as garnet and phengite, may retain pro-
grade or peak stage compositional information that has
allowed quantification of these P–T conditions as well
as those recorded during exhumation (e.g. Wei et al.,
2009, 2013; Massonne, 2011, 2012; Li et al., 2016b).
One particular challenge in modeling UHP eclogites
is the large P–T stability field of high-variance mineral
assemblages at peak conditions [e.g. Grt þ Omp þ Coe
6 Ph 6 Ky; mineral abbreviations follow Whitney &
Evans (2010)] with little change in the proportions or
compositions of the rock-forming minerals, particularly
garnet and omphacite (e.g. �Stıpska & Powell, 2005; Wei
et al., 2013; Groppo et al., 2015). Another challenge for
medium-temperature eclogites (MT eclogites) with peak
T of 550–900�C (Carswell, 1990) is the common re-
equilibration of rock-forming minerals during exhum-
ation after residence at elevated temperatures
(>700�C), which may limit our ability to determine pro-
grade, peak and retrograde P–T conditions using con-
ventional thermobarometry (Caddick et al., 2010).During the past 20 years, phase equilibrium model-
ing has extended to complex compositional systems
that are geologically realistic. To optimize the results,
the amount of ferric iron, which is commonly not ana-
lyzed or is uncertain, and an appropriate H2O content,
which is commonly less than the loss on ignition (LOI),
to be used in the bulk-rock composition must be deter-
mined; both of these may significantly influence the
phase equilibria of eclogites (e.g. Proyer et al., 2004;
Clarke et al., 2006; Rebay et al., 2010). Because these
issues have not been considered fully in recent exam-
ples of phase equilibrium modeling of MT–UHP eclo-
gites (e.g. Massonne, 2011, 2012; Wei et al., 2013;
Groppo et al., 2015), we have made a particular effort to
address them in this study of MT–UHP eclogites from
Yangkou Bay in the Sulu Belt of eastern China.
The Sulu Belt is a classic example of deep continen-
tal subduction and the UHP eclogite at Yangkou Bay is
the sole example where intergranular coesite is pre-
served in crustal eclogites (e.g. Liou & Zhang, 1996; Ye
et al., 1996; Wang et al., 2018). Despite much effort, the
peak P–T conditions and the P–T path for these eclogites
remain uncertain. For instance, using conventional ther-
mobarometry [e.g. Grt–Cpx thermometer after Holland
(1983); Grt–Cpx thermometer after Krogh Ravna (2000);
Grt–Cpx–Ph thermometer after Krogh Ravna & Terry
(2004)], P–T conditions for the peak stage were esti-
mated to be 2�9–3�4 GPa at 560–760�C (Liu et al., 2009a)
or 2�8–4�5 GPa at 800–890�C (Liou & Zhang, 1996; Zhang
& Liou, 1997; Wang et al., 2014). By contrast, reintegra-
tion of exsolved minerals in host garnet in lenses of
eclogite that occur within peridotite has suggested a
much higher pressure of >7 GPa at the peak stage (Ye
et al., 2000). Therefore, one aim of our study is to assess
whether the peak pressure that is retrieved from phase
equilibrium modeling is higher than that retrieved by
conventional thermobarometry and potentially closer to
the 7 GPa derived from reintegration of majoritic garnet.
In addition, microstructures attributed to low-volume
partial melting have been observed in the eclogite and
associated country rock gneisses from the Sulu Belt.
Partial melting in eclogites has commonly been inter-
preted as the result of breakdown of hydrous minerals
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such as white micas and/or zoisite (e.g. Zeng et al.,
2009; Zheng et al., 2011; Li et al., 2014; Wang et al.,
2014, 2016). However, primary phengite in eclogites in
these localities commonly exhibits equilibrium micro-
structures and has high Si contents that indicate forma-
tion at UHP conditions—features that suggest stability
during exhumation (e.g. Chen et al., 2014; Wang et al.,
2014). Alternatively, partial melting might be due to the
exsolution of molecular H2O and structural hydroxyl
from nominally anhydrous minerals (e.g. garnet and
omphacite) during decompression (e.g. Chen et al.,
2013, 2014; Wang et al., 2017). Therefore, another aim
of our study is to assess the extent to which melt might
have been present during the later stages of exhum-
ation of the eclogites at the Yangkou locality.
In this contribution, we use phase equilibrium model-
ing combined with average P calculations and mineral
thermobarometry to quantify the P–T conditions for four
eclogites, reconstruct a common P–T path for this locality
from these data, and discuss melting and melt crystal-
lization. The sample set comprises one bimineralic, one
kyanite-bearing and two phengite-bearing eclogites,
which are the types of eclogite that occur most common-
ly in subduction settings elsewhere, potentially making
our study of wider applicability. The P–T conditions of
such rocks have important geodynamic implications, as
discussed by Hacker (2006) and Brown & Johnson
(2018). The results of our study also provide further
understanding of melting during exhumation.
GEOLOGICAL BACKGROUND
Located in eastern China, the Sulu Belt marks the colli-
sion zone between the Yangtze and North China
Cratons. The belt is generally subdivided into a UHP
sector in the center and north, and an HP sector in the
south (Fig. 1a). Lithologies in the UHP sector comprise
dominantly orthogneiss and paragneiss, with minor
eclogite, garnet peridotite, quartzite and marble (e.g.
Zhang et al., 2009; Zheng, 2009), whereas rocks in the
HP sector comprise mainly kyanite- and topaz-bearing
quartzite, marble, paragneiss and granite gneiss, with
rare blueschist (Zhang et al., 1995).
Eclogite occurs mainly as blocks and lenses within
the gneisses, garnet peridotite and marble, and has
variable mineral assemblages including Grt þ Omp 6
Ph 6 Ky 6 Ep/Zo 6 Coe/Qz. Coesite and/or pseudo-
morphs after coesite are extensively recognized as
inclusions in rock-forming minerals in eclogite and in
zircons from various lithologies, including gneisses,
kyanite-bearing quartzite and marble (e.g. Wang et al.,
1993; Liou & Zhang, 1996; Nakamura & Hirajima, 2000;
Ye et al., 2000; Zhang et al., 2005; Liu et al., 2006b; Liu &
Liou, 2011; Wang et al., 2014). Diamonds have been
suggested to be present as inclusions in garnet (Xu
et al., 2003, 2005), as single grains in heavy mineral con-
centrates derived from eclogite (e.g. Zhang et al., 2007).
Calculated pressures and temperatures vary widely,
with estimates of pressure for peak metamorphism
from 2�8 to 4�5 GPa at temperatures ranging from
�600�C to >850�C (Liou & Zhang, 1996; Zhang & Liou,
1997; Zhang et al., 2006; Zhu et al., 2007; Liu et al.,
2009a; Liu & Liou, 2011; Wang et al., 2014). Higher peak
pressures of >7 GPa and >5�5 GPa have also been sug-
gested for unusual eclogite lenses in peridotite at
Yangkou Bay (Ye et al., 2000) and for marble at
Shanqingge (Liu et al., 2006b), respectively. Eclogites
record variable retrogression during exhumation, such
as overprinting by granulite-facies assemblages (e.g.
Wang et al., 1993; Banno et al., 2000; Nakamura &
Hirajima, 2000) or (high-pressure) amphibolite-facies
assemblages (e.g. Zhang et al., 2005; Zhu et al., 2007;
Liu et al., 2009a; Nakamura & Hirajima, 2010).
Supercritical fluid or hydrous melt may have been pre-
sent in eclogites and the associated country rocks, ei-
ther at the peak stage (e.g. Ferrando et al., 2005; Zhang
et al., 2008) or generated during decompression owing
to exhumation (e.g. Zheng et al., 2011; Chen et al., 2012,
2014; Li et al., 2014; Wang et al., 2014, 2016, 2017).Zircon geochronology, summarized by Liu et al.
(2006a) and Liu & Liou (2011), reveals the timing of the
prograde, peak and retrograde stages of the meta-
morphic evolution as follows. The prograde segment of
the P–T path into the coesite stability field occurred be-
tween c. 245 and 235 Ma, the peak segment of the P–T
path in the coesite stability field lasted from c. 235 to c.
225 Ma, and the retrograde stage through the quartz
eclogite-facies to the amphibolite-facies lasted from c.
225 to c. 210 Ma.
The eclogites in this study were sampled from
Yangkou Bay in the central Sulu Belt (Fig. 1b). The litholo-
gies include metagabbro, coesite-bearing eclogite, ser-
pentinized peridotite and quartzo-feldspathic gneiss (Liou
& Zhang, 1996; Ye et al., 2000). Eclogite is enclosed in the
gneiss and garnet peridotite. All of the UHP units are cut
by lamprophyre and quartz porphyry dikes.
PETROLOGY
Four eclogites were chosen for detailed petrographic ana-
lysis and phase equilibrium modeling. The samples com-
prise two phengite-bearing eclogites (YK128-12 and YKS-
5), one kyanite-bearing eclogite (WYKK-3) and one bimi-
neralic eclogite (YK5-2). Except for YKS-5, the samples are
characterized by intergranular coesite and inclusions of
coesite in both rock-forming and accessory minerals.
PetrographyPhengite-bearing eclogitesPhengite-bearing eclogites have the mineral assem-
blage Grt þ Omp þ Ph þ Qz 6 Coe 6 Ky, with <5 vol. %
phengite in YK128-12, but >5 vol. % phengite in YKS-5.
Both rocks have a granular microstructure, but
YK128-12 is finer grained (grain size 0�3–0�5 mm) than
YKS-5 (grain size 0�4–0�7 mm).
In YK128-12, narrow veinlets (width <0�05 mm) filled
with a symplectite of Hbl þ Pl 6 Qz 6 Cpx replace
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omphacite along grain boundaries in the primary eclog-
ite (Fig. 2a and b). The primary eclogite is also cut by
wider veins (width >0�5 mm) filled with dominantly
coarse-grained (0�3–0�7 mm) Ph þ Qz (Fig. 2a and c) and
subordinate fine-grained symplectite (normally
<0�01 mm, locally 0�03–0�1 mm) of Kfs þ Pl þ Bt þ Qz
around the coarse phengites (Fig. 2c and d). In places,
cuspate K-feldspar has small dihedral angles against
coarse quartz and phengite (Fig. 2d). Symplectites of Hbl
þ Pl 6 Qz 6 Cpx also develop along the margins of the
veins (Fig. 2a and e) and clinopyroxene separates the
symplectite from quartz along the outer edges of the
veins (Fig. 2c and e). Sometimes near veinlets, primary
omphacite or omphacite inclusions in garnet show a re-
action relationship suggesting replacement by horn-
blende, plagioclase and/or quartz (Fig. 2f), indicating that
the reaction Grt þ Omp! Hbl þ Plþ Qz may occur.
In this rock type, coesite is either included in garnet
and omphacite or occurs as intergranular grains with a
number density of >14 grains in a circle of diameter
4 mm (Fig. 2g and h). Radiating quartz fibers and/or pal-
isade quartz around polycrystalline aggregates indicate
the former presence of intergranular coesite. Granular
garnet has inclusions of phengite, omphacite, rutile and
coesite/quartz, sometimes with zircon and rarely kyanite
(Fig. 2f and i). Omphacite occurs in several microstruc-
tural settings. Primary omphacite occurs as inclusions
in garnet and as one of the main rock-forming minerals
(Cpx1, Fig. 2b, e and f), whereas secondary omphacite
occurs in veinlets around garnet or primary omphacite
(Cpx2, Fig. 2b and e), in symplectites (Cpx3) with horn-
blende, plagioclase and quartz in veinlets (Fig. 2b) or
veins (Fig. 2e), and separating symplectites from coarse
quartz at the outer edge of veins (Cpx4, Fig. 2e). Coarse
rock-forming phengite is mainly euhedral and un-
altered, whereas finer-grained phengite is subhedral to
anhedral (Fig. 2h and i).
In YKS-5 garnet and omphacite (Cpx1) grains are elon-
gated parallel to the foliation defined by phengite
(Fig. 3a), and garnet contains inclusions of phengite,
omphacite (Cpx1), quartz and rutile (Fig. 3a and b).
Coesite or quartz pseudomorphs after coesite are
absent. Although not as well developed, similar to
YK128-12, veinlets with a symplectite of Hbl þ Pl 6 Qz
occur locally along omphacite grain boundaries, but the
larger veins present in YK128-12 are absent from YKS-5.
Kyanite-bearing eclogiteSample WYKK-3 is medium-grained (0�3–1�0 mm) with
a mineral assemblage of Grt (35–40 vol. %) þ Omp (40–
45%) þ Ph (10–15%) þ Ky (3–5%) þ Coe/Qz (3–5%), and
accessary rutile and ilmenite. Anhedral to subhedral
garnet and omphacite (Cpx1) are commonly elongated
and fractured. In plane-polarized light, some larger gar-
nets show an inclusion-rich area surrounded by a clear
rim (Fig. 3c). Inclusions of omphacite (Cpx1), kyanite,
coesite/quartz and phengite, together with rutile/ilmen-
ite and zircon, are common in garnet (Fig. 3c and d).
Kyanite also replaces garnet, containing fine-grained
garnet relics (<0�01 mm; Fig. 3e). Intergranular coesite
and coesite inclusions in garnet and omphacite are par-
tially or completely replaced by quartz aggregates.
Hornblende and plagioclase occur in the veins that cut
the primary eclogite (Fig. 3f).
Fig. 1. (a) Simplified geological map of the Sulu Belt in eastern China. (b) Geological map of the Yangkou locality (based on Wanget al., 2016). TB, Tarim Block; NCB, North China Block; SCB, South China Block.
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Bimineralic eclogiteThe bimineralic eclogite (YK5-2) has an equigranular
granoblastic microstructure (grain size 0�2–0�8 mm, but
mostly 0�3–0�5 mm) with a mineral assemblage of Grt
(45–50 vol. %) þ Omp (40–45%) þ Coe/Qz (5–10%), and
accessory rutile, ilmenite, apatite and zircon. Some larger
garnets show an inclusion rich-area inferred to represent
a core surrounded by a clear mantle (Fig. 3g). The inclu-
sions are mostly omphacite (Cpx1), rutile/ilmenite and
quartz/coesite, with subordinate apatite and zircon.
Fig. 2. Photomicrographs showing the mineralogy and microstructures of phengite-bearing eclogite YK128-12 under plane-polar-ized light (a, g), cross-polarized light (h, i) and in BSE images (b–f). (a) Veins and veinlets in eclogite. Light-colored minerals in veinsare mainly coarse phengite and quartz; dark-colored minerals along veinlets and at the edge of veins are mainly Hbl þ Pl 6 Qz 6Cpx3. (b) In veinlets, a symplectite of Hbl þ Pl 6 Qz 6 Cpx3 replaces primary omphacite. Cpx1 represents primary omphacite oromphacite inclusions in garnet. Cpx2 represents retrograde omphacite with higher j(o) content along garnet or primary omphacitegrains, whereas Cpx3 represents retrograde omphacite with lower j(o) content in the symplectites. (c, d) Symplectite of Kfs þ Pl þBt 6 Qz occurs around coarse phengite in veins. (e) At the edges of veins, symplectite of Hbl þ Pl 6 Qz 6 Cpx3 develops after Cpx1.Cpx4 represents clinopyroxene with low Jd content but high Ae content rimming coarse quartz in veins. (f) Adjacent to veinlets, asymplectite of Hbl þ Pl 6 Qz replaces primary omphacite and omphacite inclusions in garnet. Ky occurs as inclusions in garnet orreplacing garnet and omphacite. (g, h) Intergranular coesite between garnet and omphacite. Quartz with radiating fibers surroundscoesite. Coesite shows higher relief than quartz. (i) Garnet showing inclusions of omphacite, phengite and quartz. Coesite occurs asinclusions in primary garnet and omphacite.
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Intergranular coesite (Fig. 3h), with a number density of
up to 13 grains in a circle of diameter 4 mm, and inclu-
sions of coesite in garnet and omphacite (Cpx1) are ubi-
quitous. Most coesite grains, regardless of location, have
been retrograded to palisade quartz or mosaic quartz at
the rims. Veinlets of symplectite comprising Cpx3þ Hbl þPl 6 Qz occur along Cpx1 or Cpx2 grain boundaries
(Fig. 3i). Adjacent to these veinlets, rare kyanite together
with quartz are present replacing garnet (Fig. 3i), whereas
intergranular coesite is absent.
Metamorphic evolutionBased on petrographic observations, several stages in
the metamorphic evolution of the eclogites may be
Fig. 3. Photomicrographs showing the mineralogy and representative microstructures of phengite eclogite YKS-5 (a, b), kyanite-bearing eclogite WYKK-3 (c–f) and bimineralic eclogite YK5-2 (g–i) under cross-polarized light (a, b, g), plane-polarized light (c, f, h),and in BSE images (d, e, i). (a, b) Inclusions of omphacite, phengite, rutile and quartz in garnet, and pristine rock-forming phengiteshowing no evidence of retrograde reaction. (c) Garnet showing color differences. The inner part is inclusion-rich, interpreted torepresent the garnet core, whereas the mantle and rim are devoid of inclusions. The inclusions are mostly omphacite, phengite andquartz, with subordinate rutile, ilmenite and apatite. Rock-forming phengite is subhedral in shape and pristine. (d) Kyanite as inclu-sions in garnet. (e) Kyanite with garnet inclusions. (f) Retrogression along infilled fractures with hornblende and plagioclase in theveins. (g) Garnet showing an inclusion-rich core surrounded by a clear rim. The inclusions are mostly omphacite, rutile/ilmeniteand quartz/coesite, with subordinate apatite and zircon. (h) Intergranular coesite (showing higher relief than quartz) between garnetand omphacite. (i) Veinlets with symplectite comprising Cpx þ Hbl þ Pl 6 Qz; there are rare kyanite relics.
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inferred (Table 1). For both phengite-bearing eclogites,
YK128-12 and YKS-5, a prograde stage is evidenced by
inclusions of Cpx1, Ph, Ky, Coe and Rt in Grt. The peak
stage without Ky in the mineral assemblage (Grt, Cpx1,
Ph, Coe and Rt, Table 1) could be due to the consump-
tion of Ky during the prograde evolution. An early de-
compression stage is represented by the retrograde
mineral assemblage Cpx2, Qz, Rt and secondary Ky
replacing Cpx1, and is followed by a late decompres-
sion stage recorded in the veins and veinlets by sym-
plectites of Kfs þ Bt þ Pl þ Qz around coarse phengite,
symplectites of Hbl þ Pl 6 Qz 6 Cpx3 replacing Cpx1
and Cpx2, and the development of Cpx4 against Qz. The
Ky-bearing eclogite, WYKK-3, has a prograde mineral
assemblage similar to that in the Ph-bearing eclogite, as
indicated by inclusions of Cpx1, Ph, Ky, Coe and Rt in
Grt. The peak stage has a mineral assemblage of Grt þCpx1 þ Ph þ Coe þ Ky þ Rt. In the veins, the mineral as-
semblage of Hbl þ Pl þ Qz probably indicates a late hy-
dration stage. For the bimineralic eclogite, YK5-2, the
prograde evolution to the peak stage did not change the
mineral assemblage (Grt, Cpx1, Coe and Rt). The early
decompression stage is identified by Cpx2, Qz, Rt and
secondary Ky, and replacement of Cpx1, followed by
the late decompression stage recorded by symplectites
of Hbl þ Pl 6 Qz replacing Cpx1 and Cpx2.
Mineral compositionsAnalytical methodsMineral analyses were performed using a JEOL 8900
electron probe microanalyzer at the Advanced Imaging
and Microscopy Laboratory at the University of
Maryland. The following operating conditions were uti-
lized: 15 kV accelerating voltage, 10 nA (phengite) or
25 nA (remaining phases) cup current, and a 1–10 lm
beam diameter. A series of natural and synthetic stand-
ards was utilized: plagioclase [Lake County plagioclase
(Si, Al, Na, K), Kakanui hornblende (Fe, Mg, Mn), micro-
cline (K), Broken Hill rhodonite (Mn)]; phengite
[Muthuen Township muscovite (K, Al, Si), Tiburon
Peninsula albite (Na), Kakanui hornblende (Fe, Mg, Ti),
Geophysical Lab synthetic Ba-glass (Ba), Lake County
plagioclase (Ca), and Broken Hill Rhodonite (Mn)]; kyan-
ite [Minas Gerais kyanite (Si, Al), Kakanui hornblende
(Fe, Mg, Ti, Ca), Broken Hill rhodonite (Mn), and Tiburon
Peninsula albite (Na)]; garnet [USGS GTAL garnet (Si,
Al, Fe, Ca, Mg), Kakanui hornblende (Na, K, Ti),
Johnstown hypersthene (Cr), and Minas Gerais spes-
sartine (Mn)]; and hornblende, biotite and omphacite
[Kakanui hornblende (Si, Til, Al, Fe, Mg, Ca, K and Na),
Johnstown hypersthene (Cr), and Minas Gerais spes-
sartine (Mn)]. Raw X-ray intensities were corrected
using a ZAF algorithm.
GarnetIn the phengite-bearing eclogite YK128-12, element
mapping of one of the larger garnets, supported by a
traverse of point analyses, shows relict zoning in cal-
cium, but only weak or no zoning in other cations
(Fig. 4a and b). In the apparent core the grossular con-
tent is relatively low (29 mol %), increasing to 33 mol %
in the mantle before decreasing outwards to 30 mol %
at the edge; there is an increase to 32 mol % in some
parts of the outermost rim or adjacent to fractures
(Table 2); XFe [¼ Fe/(Fe þ Mg)] changes little from cen-
ter to edge (0�67–0�70). For the phengite-bearing eclog-
ite YKS-5, calcium is seen to be zoned (Fig. 4c and d);
the grossular content in the apparent core is 29 mol %,
which decreases to 25 mol % in the mantle. The highest
grossular content (33 mol %) develops in the outmost
rim or adjacent to fractures (Fig. 4c; Table 2). XFe
changes little from center to edge (0�59–0�61).
For the kyanite-bearing eclogite (WYKK-3), element
mapping of a large garnet with an inclusion-rich appar-
ent core and a clear mantle/rim, supported by a traverse
of point analyses, shows relict zoning in calcium but
less distinct zoning of magnesium and iron (Fig. 4e and
f). In one half of the garnet (from core to A in Fig. 4f), the
grossular content in the core is relatively high (31–
32 mol %), decreasing to 26 mol % in the inner rim and
then increasing to 34 mol % in the outer rim (Table 2),
resembling the zoning pattern in the garnet from one of
the phengite-bearing eclogites (YKS-5); XFe varies little
Table 1: Metamorphic evolution inferred for the eclogites at Yangkou Bay, Sulu Belt
Prograde stage Peak stage Early decompression stage Late decompression stage
Samples YK128-12 and YKS-5Grt and inclusions
(Cpx1, Ph, Ky, Coe, Rt)Grt, Cpx1, Ph, Coe, Rt Ky (replacing Cpx1),
Cpx2, (Ph), Qz, Rt(4a) Hbl, Pl, Qz, (Cpx3) (symplectite
around Cpx1 and Cpx2)(4b) Kfs, Bt, Pl, Qz (between Ph and Qz)(4c) Cpx4 around Qz
Prograde stage Peak stage Early decompression stage Late hydration stage
Sample WYKK-3Grt and inclusions
(Cpx1, Ph, Ky, Coe, Rt)Grt, Cpx1, Ph, Ky, Coe, Rt Ky (replacing Grt þ Cpx1),
Cpx2, (Ph), Qz, RtHbl, Pl, Qz (in the fractures)
Prograde stage Peak stage Early decompression stage Late decompression stage
Sample YK5-2Grt and inclusions
(Cpx1, Coe, Rt)Grt, Cpx1, Coe, Rt Ky (replacing Cpx1),
Cpx2, Qz, RtHbl, Pl, Qz, (Cpx3) (symplectite
around Cpx1 and Cpx2)
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Fig. 4. Ca element maps and traverses of point analyses for selected garnets from the four studied eclogites. Ca content in garnetincreases from green to yellow to red in the images (see online version for color images).
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(0�64–0�66) from core to rim. In the other half of the gar-
net (from core to B in Fig. 4f), the grossular content
increases from core (30–31 mol %) to rim (34 mol %),
with XFe of 0�65–0�68.
In the bimineralic eclogite (YK5-2), a larger garnet
shows complex compositional zoning from an inclusion-
rich apparent core across the associated mantle (Fig. 4g
and h). The traverse shows that in one half of the garnet,
the grossular content decreases from an inner core (31–
32 mol %) to an outer core (28–29 mol %), before increas-
ing to 33 mol % in the mantle and then decreasing to
29 mol % in the inner rim (Table 2). In the outer rim, the
grossular content increases to 32 mol % again. The other
half of the garnet shows similar grossular variations.
Detailed examination of the core shows that inclusions
are rare in the inner part with the higher grossular con-
tent, whereas the outer part with the lower grossular con-
tent has numerous omphacite inclusions. Although care
was taken during analysis to avoid inclusions, the low
grossular content could reflect contamination of the anal-
yses in this area by cryptic fine-grained inclusions of
omphacite. Overall, the grossular zoning is similar to that
in garnet from one of the phengite-bearing eclogites
(YK128-12). The almandine and pyrope contents are simi-
lar to those of the other samples, but there is no systemat-
ic zoning, and XFe varies only slightly (0�67–0�71).
In summary, garnet in all four eclogites retains relict
zoning in calcium content whereas other bivalent cati-
ons preserve only weak zoning or are unzoned. Two
contrasting calcium zoning profiles are present. The
first shows an increase, followed by a decrease and
then an increase in calcium from the core to the rim, as
recorded by the larger garnets in YK128-12 and YK5-2.
The second shows a decrease followed by an increase
in calcium from core to rim or against fractures, as
recorded by the larger garnets in YKS-5 and WYKK-3.
OmphaciteWithin each sample, the primary omphacite, whether as
inclusions in garnet or as a main rock-forming mineral,
has indistinguishable compositions (Table 3, Fig. 5a).
The j(o) [where j(o) ¼ Jd þ Ae] content of the rock-
forming omphacite generally shows no systematic var-
iations from core to rim.For the phengite-bearing eclogite YK128-12, ompha-
cite inclusions (cpx1) in garnet have j(o) of 0�62 and
rock-forming omphacite (Cpx1) has j(o) of 0�60–0�64.
Secondary omphacite (Cpx2) in veinlets around garnet
or primary omphacite has slightly lower j(o) of 0�52–
0�57, whereas sodic clinopyroxene in symplectites in
veinlets (Cpx3) or against coarse quartz in veins (Cpx4)
has a significantly lower j(o) of 0�26–0�28 and 0�24–0�27,
Table 2: Representative compositions of garnet from the eclogites at Yangkou, Sulu Belt
YK128-12 YKS-5 WYKK-3 YK5-2
Position: core core mantle rim core mantle mantle rim core core mantle rim core core mantle rim
SiO2 39�10 38�88 38�61 38�87 39�26 38�99 39�35 39�13 38�95 38�84 38�99 39�64 38�51 38�60 38�83 38�45TiO2 b.d. b.d. 0�04 0�03 b.d. b.d. 0�03 b.d. 0�07 0�09 b.d. b.d. 0�03 0�06 b.d. 0�05Al2O3 21�89 21�54 21�72 21�81 22�03 22�18 22�29 22�19 21�28 21�34 21�74 22�32 21�52 21�30 21�63 21�58Cr2O3 b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. 0�05 b.d. b.d. b.d. b.d. b.d. b.d. b.d.FeO 22�24 21�90 21�29 21�07 20�00 20�89 20�47 19�27 21�51 21�49 21�77 19�67 21�93 22�11 22�15 21�73MnO 0�46 0�41 0�45 0�41 0�40 0�50 0�46 0�44 0�51 0�50 0�59 0�46 0�40 0�43 0�44 0�43MgO 6�05 5�79 5�53 5�59 7�31 7�95 7�73 7�02 5�94 6�12 6�85 6�07 5�65 5�68 5�82 5�55CaO 10�52 10�96 11�90 11�67 10�63 9�15 9�96 11�35 11�55 11�34 9�38 12�34 11�45 11�32 10�69 11�59Na2O b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.K2O b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.Total 100�26 99�48 9�54 99�45 99�63 99�69 100�28 99�38 99�81 99�72 99�31 100�49 99�48 99�49 99�58 99�37Oxygens 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00 12�00Si 3�00 3�01 2�99 3�01 3�00 2�98 2�99 2�99 3�00 3�00 3�01 3�01 2�98 2�99 3�01 2�98Ti b.d. b.d. 0�00 0�00 b.d. b.d. 0�00 0�00 0�00 0�01 0�00 0�00 0�00 0�00 0�00 0�00Al 1�98 1�97 1�98 1�99 1�99 2�00 2�00 2�00 1�93 1�94 1�98 2�00 1�97 1�95 1�97 1�97Cr b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. 0�00 b.d. b.d. b.d. b.d. b.d. b.d. b.d.Fe3þ 0�01 0�01 0�04 0�00 0�01 0�04 0�03 0�02 0�05 0�06 0�00 0�00 0�07 0�06 0�01 0�06Fe2þ 1�42 1�41 1�34 1�36 1�27 1�29 1�28 1�23 1�34 1�33 1�41 1�25 1�35 1�37 1�42 1�35Mn 0�03 0�03 0�03 0�03 0�03 0�03 0�03 0�03 0�03 0�03 0�04 0�03 0�03 0�03 0�03 0�03Mg 0�69 0�67 0�64 0�64 0�83 0�91 0�88 0�79 0�68 0�70 0�79 0�69 0�65 0�66 0�67 0�64Ca 0�87 0�91 0�99 0�97 0�87 0�75 0�81 0�95 0�95 0�94 0�78 1�01 0�95 0�94 0�89 0�96Na b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.K b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.Sum 8�00 8�00 8�00 8�00 8�00 8�00 8�00 8�00 8�00 8�00 8�00 7�99 8�00 8�00 8�00 8�00Alm 0�47 0�47 0�45 0�45 0�42 0�43 0�43 0�41 0�44 0�44 0�47 0�42 0�45 0�46 0�47 0�45Grs 0�29 0�30 0�33 0�32 0�29 0�25 0�27 0�32 0�32 0�31 0�26 0�34 0�32 0�31 0�29 0�32Pyr 0�23 0�22 0�21 0�21 0�28 0�30 0�29 0�26 0�23 0�23 0�26 0�23 0�22 0�22 0�22 0�21Sps 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01 0�01Fe# 0�67 0�68 0�68 0�68 0�60 0�59 0�59 0�61 0�66 0�65 0�64 0�64 0�67 0�68 0�68 0�68
Fe# ¼ Fe2þ/(Fe2þ þ Mg); Alm, almandine; Grs, grossular; Spss, spessartine; Pyr, pyrope. The mineral formulae were calculatedusing the program AX (T. Holland’s AX software page: http://www.esc.cam.ac.uk/research/research-groups/research-projects/tim-hollands-software-pages/ax). b.d., below detection.
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Fig. 5. Compositions of omphacite and hornblende in the eclogites. (a) Ternary classification diagram for sodic clinopyroxene.Black filled symbols represent omphacite inclusions in garnet (Cpx1) and open symbols represent primary rock-forming omphacite(Cpx1). (b, c) Compositions of hornblende (after Leake et al., 1997).
Table 3: Representative compositions of omphacite and hornblende from the eclogites at Yangkou, Sulu Belt
YK128-12 YKS-5 WYKK-3 YK5-2
Cpx1 Cpx1 Cpx2 Cpx2 Cpx3 Cpx4 Hbl Cpx1 Cpx1 Cpx1 Cpx1 Hbl Cpx1 Cpx1 Hbl Hbl
SiO2 56�93 56�91 56�05 54�83 52�99 52�97 44�11 56�45 56�24 56�54 56�63 49�95 56�55 56�14 46�67 43�47TiO2 0�05 0�03 0�04 0�06 0�13 0�04 0�18 0�06 b.d. 0�04 b.d. 0�29 0�04 0�04 0�13 0�55Al2O3 15�48 15�60 13�55 12�50 6�87 2�37 14�85 13�35 12�19 14�53 13�64 7�20 13�81 15�39 10�87 13�60Cr2O3 b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.FeO 4�10 3�62 4�89 5�84 6�14 12�59 11�06 3�52 3�06 4�06 3�52 11�45 3�89 3�82 9�99 13�49MnO b.d. b.d. b.d. 0�03 b.d. 0�09 0�05 0�04 0�05 b.d. 0�03 0�20 b.d. b.d. 0�05 0�07MgO 5�94 5�58 6�91 7�29 11�82 10�06 12�46 7�69 8�24 6�44 7�24 15�40 7�04 5�83 15�45 12�49CaO 9�46 8�88 11�05 12�63 18�72 18�34 11�98 11�91 12�54 10�19 11�06 11�33 10�77 9�57 11�64 11�20Na2O 9�01 9�18 8�21 7�22 3�41 3�62 3�68 7�66 7�22 8�61 8�00 1�79 8�26 8�81 2�32 2�98K2O b.d. b.d. b.d. b.d. b.d. b.d. 0�06 b.d. b.d. b.d. b.d. 0�22 b.d. b.d. 0�08 0�04Total 100�97 99�79 100�70 100�41 100�07 100�09 98�37 100�67 99�55 100�42 100�11 97�61 100�37 99�61 97�12 97�86Oxygens 6�00 6�00 6�00 6�00 6�00 6�00 23�00 6�00 6�00 6�00 6�00 23�00 6�00 6�00 23�00 23�00Si 1�98 1�99 1�96 1�94 1�92 1�96 6�38 1�98 1�99 1�98 1�99 7�15 1�98 1�98 6�71 6�33Ti 0�00 0�00 0�00 0�00 0�00 0�00 0�02 0�00 b.d. 0�00 b.d. 0�03 0�00 0�00 0�02 0�06Al 0�63 0�65 0�56 0�52 0�29 0�10 2�53 0�55 0�51 0�60 0�57 1�22 0�57 0�64 1�84 2�34Cr b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.Fe3þ 0�01 0�00 0�07 0�09 0�11 0�23 0�00 0�02 0�00 0�03 0�00 0�26 0�03 0�01 0�29 0�34Fe2þ 0�11 0�11 0�07 0�09 0�08 0�16 1�34 0�09 0�09 0�09 0�10 1�11 0�09 0�11 0�91 1�31Mn b.d. b.d. b.d. 0�00 b.d. 0�00 0�01 0�00 0�00 b.d. 0�00 0�02 b.d. b.d. 0�01 0�01Mg 0�31 0�29 0�36 0�39 0�64 0�56 2�69 0�40 0�44 0�34 0�38 3�29 0�37 0�31 3�31 2�71Ca 0�35 0�33 0�42 0�48 0�73 0�73 1�86 0�45 0�48 0�38 0�42 1�74 0�40 0�36 1�79 1�75Na 0�61 0�62 0�56 0�50 0�24 0�26 1�03 0�52 0�50 0�58 0�55 0�50 0�56 0�60 0�65 0�84K b.d. b.d. b.d. b.d. b.d. b.d. 0�01 b.d. b.d. b.d. b.d. 0�04 b.d. b.d. 0�02 0�01Sum 4�00 3�99 4�00 4�00 4�00 4�00 15�86 4�00 4�00 4�00 4�00 15�44 4�00 4�00 15�64 15�81Jd 0�62 0�64 0�54 0�49 0�23 0�07 0�54 0�50 0�59 0�56 0�56 0�62Ae 0�00 0�00 0�03 0�03 0�03 0�20 0�00 0�00 0�00 0�00 0�01 0�00WEF 0�38 0�36 0�42 0�48 0�75 0�73 0�46 0�50 0�40 0�44 0�43 0�38Jo¼JdþAe
0�62 0�64 0�57 0�52 0�26 0�27 0�54 0�50 0�59 0�56 0�57 0�62
Jd, jadeite; Ae, aegirine; WEF, wollastonite þ enstatite þ ferrosilite. Other details are as in the footnote to Table 2.
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respectively. However, Cpx4 is rich in the Ae compo-
nent (0�16–0�20), whereas its Jd content is only 0�07–
0�08, whereas Cpx1–Cpx3 are rich in the Jd component
with a low Ae of <0�04 (Fig. 5a). For the other three eclo-
gites, omphacite as inclusions in garnet and rock-
forming omphacite (Cpx1) has j(o) of 0�51–0�56 and
0�50–0�54 (phengite-bearing eclogite YKS-5), 0�51–0�59
and 0�53–0�56 (kyanite-bearing eclogite WYKK-3) and
0�57–0�60 and 0�56–0�63 (bimineralic eclogite YK5-2),
respectively.
PhengiteThe compositional variation of phengite from different
microstructural settings is shown in Supplementary
Data Electronic Appendix Fig. S1 (supplementary data
are available for downloading at http://www.petrology.
oxfordjournals.org). For YK128-12, phengite included in
garnet has TiO2 contents from 0�48 to 0�58 wt % and Si
contents from 3�48 to 3�54 p.f.u. (per formula unit; 11 O
basis). Coarse phengite flakes have TiO2 contents of
0�34–0�62 wt % and Si contents of 3�49–3�54 p.f.u. Rarer
fine-grained phengite has TiO2 contents of 0�54–0�71 wt
% and Si contents of 3�43–3�44 p.f.u., slightly lower than
the other types. The large phengite flakes (>0�7 mm) in
the vein have TiO2 contents of 0�34–0�68 wt % and Si
contents of 3�44–3�53 p.f.u., similar to the coarse phen-
gite and the inclusions in garnet. For YKS-5 and WYKK-
3, phengite included in garnet has TiO2 contents of
0�61 and 0�42–0�49 wt %, and Si contents of �3�51 and
3�47–3�53 p.f.u., respectively, whereas phengite outside
garnet has TiO2 contents of 0�37–0�61 and 0�45–0�62 wt
%, and Si contents of 3�47–3�58 and 3�42–3�53 p.f.u,
respectively. BaO contents in all phengite are in the
range from 0�02 to 0�71 wt % (Table 4).
HornblendeHornblende in the symplectites has Al2O3 contents of
10�87–17�98 wt %. It is pargasite (Fig. 5c; Leake et al.,
1997) in YK128-12, with Mg# [Mg/(Mg þ Fe2þ)] of 0�55–
0�72 and (Na þ K)A from 0�85 to 0�94 p.f.u. (23 O basis),
whereas it ranges from edenite to pargasite (Fig. 5c) in
YK5-2, with Mg# of 0�66–0�78 and (Na þ K)A of 0�54–
0�80 p.f.u. Hornblende bordering the veins in the
kyanite-bearing eclogite (WYKK-3) has lower Al2O3 con-
tents (5�09–8�50 wt %) than the hornblende in YK128-12
and YK5-2. It is magnesiohornblende (Fig. 5b; Leake
et al., 1997), with Mg# of 0�72–0�79 and (Na þ K)A of
0�28–0�42 p.f.u.
Other mineralsThe sodic feldspar in the symplectites in the two
phengite-bearing eclogites (YK128-12 and YK5-2) is alb-
ite with Ab0�91–0�98, whereas the plagioclase in the veins
in the kyanite-bearing eclogite (WYKK-3) is oligoclase
with Ab0�81–0�90 (Table 4). Kyanite in all samples is nearly
pure Al2SiO5 (Table 4).
PHASE EQUILIBRIUM MODELING
The purpose of the phase equilibrium modeling is to
quantify the P–T evolution of the eclogites at Yangkou
Bay in the Sulu Belt.
Table 4: Representative composition of phengite, biotite, plagioclase and kyanite from the eclogites at Yangkou, Sulu Belt
YK128-12 YKS-5 WYKK-3
Mineral:. Phflake
Phtiny
Ph inGrt
Ph inveins
Ph inveins
Bt inveins
Bt inveins
Pl inveins
Phflake
Ph inGrt
Phflake
Ph inGrt
Pl Ky inGrt
Kyout
SiO2 53�02 51�98 53�72 51�83 53�11 39�99 39�01 66�36 54�06 52�02 53�32 53�46 64�58 36�87 37�31TiO2 0�34 0�54 0�50 0�68 0�34 2�35 2�86 b.d. 0�42 0�61 0�48 0�47 0�04 b.d. b.d.Al2O3 24�31 26�56 24�25 26�43 24�40 17�22 16�69 20�62 23�47 24�50 24�63 24�85 20�32 62�47 63�08FeO 1�73 2�06 1�84 2�07 1�72 11�92 12�65 0�42 1�52 1�43 1�62 1�57 0�81 0�48 0�54MnO b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.MgO 4�47 3�64 4�64 3�55 4�50 15�24 13�91 b.d. 5�02 4�28 4�55 4�46 0�96 b.d. b.d.CaO b.d. 0�04 b.d. b.d. b.d. 0�06 b.d. 1�72 b.d. b.d. b.d. b.d. 2�86 b.d. b.d.Na2O 0�15 0�57 0�29 0�19 0�12 0�43 0�11 10�81 0�11 0�27 0�16 0�21 9�63 b.d. b.d.K2O 10�60 10�02 10�68 10�92 11�03 9�86 9�67 0�04 10�70 10�51 10�54 10�53 0�07 b.d. b.d.BaO 0�49 0�67 0�32 0�23 0�31 0�19 0�46 b.d. 0�14 0�67 0�34 0�56 b.d. b.d. b.d.Total 95�12 96�08 96�23 95�88 95�53 97�25 95�36 99�96 95�44 94�30 95�64 96�11 99�28 99�82 100�93O 11�00 11�00 11�00 11�00 11�00 11�00 11�00 8�00 11�00 11�00 11�00 11�00 8�00 5�00 5�00Si 3�54 3�44 3�54 3�44 3�53 2�87 2�87 2�92 3�58 3�51 3�53 3�53 2�87 1�00 1�00Ti 0�02 0�03 0�03 0�03 0�02 0�13 0�16 b.d. 0�02 0�03 0�02 0�02 0�00 b.d. b.d.Al 1�91 2�07 1�89 2�07 1�91 1�45 1�45 1�07 1�83 1�95 1�92 1�93 1�07 1�99 1�99Fe2þ 0�10 0�11 0�10 0�11 0�10 0�71 0�78 0�02 0�08 0�08 0�09 0�09 0�03 0�01 0�01Mn b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d. b.d.Mg 0�44 0�36 0�46 0�35 0�45 1�63 1�52 0�00 0�50 0�43 0�45 0�44 0�06 b.d. b.d.Ca b.d. 0�00 b.d. b.d. b.d. 0�00 b.d. 0�08 b.d. b.d. b.d. b.d. 0�14 b.d. b.d.Na 0�02 0�07 0�04 0�02 0�02 0�06 0�02 0�92 0�02 0�04 0�02 0�03 0�83 b.d. b.d.K 0�90 0�85 0�90 0�92 0�94 0�90 0�91 0�00 0�91 0�90 0�89 0�89 0�00 b.d. b.d.Ba 0�01 0�02 0�01 0�01 0�01 0�01 0�01 b.d. 0�00 0�02 0�01 0�01 b.d. b.d. b.d.Total 6�95 6�96 6�96 6�97 6�97 7�76 7�71 5�01 6�94 6�96 6�94 6�94 5�01 3�00 3�00
The details are the same as in the footnote to Table 2.
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MethodsPhase diagrams were calculated using the
THERMOCALC software (Powell & Holland, 1988; ver-
sion 3.40, released on 20 March 2014) and the associ-
ated internally consistent thermodynamic dataset ds62
(Holland & Powell, 2011; updated at 6 February 2012)
for compositions in the NC(K)FMASHTO [Na2O–CaO(–
K2O)–FeO–MgO–Al2O3–SiO2–H2O–TiO2–O] system.
There are no potassic minerals in the bimineralic eclog-
ite (YK5-2), so K2O was not included in the modeling; in
the natural sample, the trivial amount of K2O is
assumed to be incorporated into pyroxene and horn-
blende. A–x relationships used in the modeling are as
follows: clinopyroxene (Green et al., 2016); garnet,
orthopyroxene and phengitic muscovite (White et al.,
2014); paragonite (Coggon & Holland, 2002); K-feldspar
and plagioclase (Holland & Powell, 2003). Pure phases
include lawsonite, kyanite, quartz/coesite, rutile, talc
and aqueous fluid (H2O). The modeling to constrain the
late prograde, peak and early retrograde P–T evolution
was done without melt as a phase, because there is no
appropriate activity–composition (a–x) model for melt
in basic rocks at these ultrahigh pressures (see Green
et al., 2016).
The bulk-rock compositions of YK5-2 and YKS-5
were obtained by inductively coupled plasma optical
emission spectrometry. For the modeling, the CaO and
FeO contents were corrected for the P2O5 and TiO2 con-
tained in apatite and ilmenite, respectively. For samples
YK128-12 and WYKK-3, the bulk-rock compositions
were considered to be inappropriate for modeling
owing to the scale of heterogeneity in these two sam-
ples. Instead, an effective domainal composition for the
primary eclogite away from the veinlets and veins was
calculated by integrating mineral compositions and vol-
umes. Primary sample compositions for YKS-5 and
YK5-2 (in wt%) together with all modified compositions
used to calculate phase diagrams (in mol %) are given
in Table 5.
H2O-present versus H2O-absent metamorphismIn their phase equilibrium modeling of MT–UHP eclo-
gites from the south Dabie orogen, a continuation to the
west of the same continental collisional event that
formed the Sulu Belt, Wei et al. (2013) assumed H2O
was present in excess. Thus, we begin the phase equi-
librium modeling by calculating a P–T pseudosection
with H2O in excess for one of the phengite-bearing eclo-
gites (YK128-12). This exercise also provides a founda-
tion for making decisions with respect to the T/P–O and
T/P–H2O modeling that follows.
The P–T pseudosection in Fig. 6 was calculated for a
range of P¼ 1�5–5�5 GPa and T¼ 600–900�C in the
NCKFMASHTO system, using the modified composition
given in Table 5 (for Fig. 6) with excess H2O and an
assumed value of 1�14 mol % for O, which is the aver-
age for eclogites retrieved from the Chinese Continental
Scientific Drilling (CCSD) main hole in the southern
Table 5: Primary sample compositions (in wt %) together with modified bulk-rock compositions used for phase diagram calculation(in mol %) for eclogites at Yangkou, Sulu Belt
Sample Figures H2O SiO2 Al2O3 CaO MgO FeOT K2O Na2O TiO2 O
YKS-5 (wt %) * 50�09 17�01 10�48 6�58 10�33 0�74 3�39 1�38 *YK5-2 (wt %) * 49�97 17�33 10�13 5�46 11�38 0�12 4�05 1�55 *YK128-12 (mol %) 6, P–T excess 52�94 10�74 10�93 7�93 9�87 0�13 4�44 1�87 1�14
7, T/P–O excess 53�55 10�87 11�06 8�02 9�99 0�13 4�49 1�89 0�0051�95 10�54 10�73 7�78 9�69 0�13 4�35 1�83 3�00
8, T–H2O 0�00 53�42 10�84 11�03 8�00 9�96 0�13 4�48 1�89 0�251�00 52�88 10�73 10�92 7�92 9�86 0�13 4�43 1�87 0�25
9, P–T 0�26 53�28 10�81 11�00 7�98 9�94 0�13 4�46 1�88 0�2511, P–T 7�15 76�62 7�96 0 3�17 0�81 3�70 0 0�19 0�40S3a, T–O 0�26 53�41 10�84 11�03 8�00 9�96 0�13 4�48 1�89 0�00S3b, P–O 0�26 52�34 10�62 10�81 7�84 9�76 0�13 4�39 1�85 2�00S4, T–O 7�17 76�93 7�99 0�00 3�18 0�82 3�71 0�00 0�20 0�00
7�10 76�16 7�91 0�00 3�15 0�81 3�67 0�00 0�19 1�00YKS-5 (mol %) S1a, T–O excess 53�18 10�64 11�51 10�42 9�17 0�50 3�48 1�11 0�00
51�58 10�32 11�16 10�11 8�89 0�48 3�38 1�07 3�00S2a, T–H2O 0 52�96 10�60 11�46 10�38 9�13 0�50 3�47 1�10 0�40
6�00 48�28 9�66 10�45 9�46 8�32 0�45 3�16 1�00 0�4010a, P–T 0�96 52�45 10�50 11�35 10�28 9�04 0�49 3�44 1�09 0�40
WYKK-3 (mol %) S1b, T–O excess 53�36 12�48 10�71 9�08 8�75 0�75 3�79 1�07 0�0051�76 12�10 10�39 8�81 8�49 0�73 3�67 1�04 3�00
S2b, T–H2O 0�00 53�20 12�44 10�68 9�05 8�73 0�75 3�77 1�07 0�3010�00 47�87 11�19 9�61 8�14 7�85 0�68 3�40 0�96 0�30
10b, P–T 1�45 52�43 12�26 10�53 8�92 8�60 0�74 3�72 1�05 0�30YK5-2 (mol %) S1c, T–O excess 53�56 10�94 11�11 8�72 10�20 0�00 4�21 1�25 0�00
51�95 10�62 10�78 8�46 9�89 0�00 4�09 1�22 3�0010c, P–T 0�00 53�34 10�90 11�07 8�69 10�16 0�00 4�20 1�25 0�40
FeOT represents total iron. O (in mol %) is equal to Fe2O3 (in mol %). Figures S1a–c, S2a, b, S3a, b and S4 are Supplementary DataElectronic Appendix at Journal of Petrology online.* Represents undetected elements in the bulk rock composition.
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Sulu Belt, �280 km south of Yangkou Bay (Zhang et al.,
2006). This pseudosection is dominated by a small
number of high-variance phase assemblage fields; inparticular, over a range of pressure, Grt þ Omp þ Law
þ Coe þ Ph þ H2O is stable at lower temperature and
Grt þ Omp þ Ky þ Coe þ Ph þ H2O is stable at higher
temperature. Wei et al. (2013) suggested using the high-
est Si content of phengite and the highest grossular
content in garnet to constrain close to peak P–T condi-
tions, although the amount of inevitable re-equilibrationof prograde zoning in garnet while reaching peak T is
unknown (Caddick et al., 2010).
For YK128-12, isopleths for the highest Si content of
phengite (3�54 p.f.u.) and the highest grossular content
in garnet (0�33) intersect in the phase assemblage field
Grt þ Omp þ Law þ Ph þ Coe þ H2O, yielding apparentP–T conditions of �4�5 GPa at �800�C. Calculated min-
eral mole proportions (on a one-oxide basis, approxi-
mating vol. %) at this P–T condition are Grt ¼ 38�4%,
Omp ¼ 42�3%, Law ¼ 9�1%, Ph ¼ 1�8%, Coe ¼ 6�5% and
Rt ¼ 1�9 %. This mineral assemblage includes �9 mol %
lawsonite, but lawsonite (or pseudomorphs after law-
sonite) is not present in the samples of this study andhas not been reported previously in eclogites from the
Yangkou locality or from the Sulu Belt. Thus, we con-
clude that the assumption of H2O present in excess is
inappropriate for these eclogites.
In support of this conclusion, we note also that at
Yangkou Bay, the presence of intergranular coesite inthe eclogites and the preservation of igneous mineral-
ogy and texture in the protolith gabbro and the
surrounding gneisses (Liou & Zhang, 1996; Wallis et al.,
1997; Zhang & Liou, 1997; Wang et al., 2018) indicate
that at the metamorphic peak these rocks were very
probably fluid-absent (Mosenfelder et al., 2005; Young
& Kylander-Clark, 2015). In addition, during eclogite-fa-
cies metamorphism both the stability of lawsonite and
the proportion that may be present are determined by
the amount of H2O in the system and the oxidation con-
ditions during metamorphism (Rebay et al., 2010).
Thus, before calculating the remainder of the P–T pseu-
dosections used in this study, we followed an iterative
process to estimate appropriate values for the ferric
iron (O) and H2O contents in the bulk compositions to
be modeled.
Estimation of Fe31 (O)To evaluate the influence of O on phase relations and
mineral stability, T–O and P–O pseudosections were cal-
culated for YK128-12 using the bulk-rock composition
listed in Table 5 (for Fig. 7). For these calculations, H2O
was assumed to be present in excess. We considered a
range of O contents from 0 to 3 mol %, where zero rep-
resents all Fe as Fe2þ and 3 represents 58�9 mol % Fe as
Fe3þ. The T–O diagram was calculated at P¼ 4�0 GPa
and the P–O diagram was calculated at T¼800�C. Using
slightly different values for P or T in modeling these T/
P–O diagrams has little influence on the results.The calculated T–O diagram is dominated by four-
and five-variant fields (Fig. 7a). Lawsonite-bearing
phase assemblages develop at T<790�C and Ky-
bearing phase assemblages develop at T> 720�C, at
O> 0�6 mol %. The mineral assemblage observed in the
sample (Grt þ Omp þ Coe þ Ph þ Rt) is present in a six-
variant phase assemblage field where O< 0�60 mol %
and T is above �725�C. The P–O pseudosection calcu-
lated at T¼ 800�C yields a similar result (Fig. 7b). Thus,
any arbitrary value for O that is <0�60 mol % in the bulk-
rock composition can be used for further modeling. As
a further constraint, we used the mineral modes and
compositions (with Fe3þ �0�03 p.f.u. for both garnet and
omphacite) to calculate an O content, which gives a
value of about 0�2 mol %. Given the low Fe3þ content of
the rock-forming minerals in this sample (Cpx with
Acm0–3 and Grt with Adr0–2), a low O content of 0�25 mol
% in the bulk-rock composition was used for further
modeling.
Following the same procedure, T–O pseudosections
were calculated for the other three samples
(Supplementary Data Electronic Appendix Fig. S2a–c).
For the phengite-bearing eclogite, YKS-5, the modeled
mineral assemblage of Grt þ Omp þ Ph þ Coe þ Rt
occurs in a field with an O content of <0�8 mol %
(Supplementary Data Electronic Appendix Fig. S2a). For
the bimineralic eclogite, YK5-2, the T–O pseudosection
was calculated in the K-absent system. Nonetheless,
the phase relations are similar to those for the phengite-
bearing eclogites (Supplementary Data Electronic
Appendix Fig. S2c). The mineral assemblage of Grt þ
Fig. 6. P–T pseudosection for phengite-bearing eclogite YK128-12. The bulk-rock composition is given in Table 5. ‘cg’ repre-sents a calculated isopleth for the highest value of grossularcontent in garnet and ‘Si’ represents a calculated isopleth forthe highest value of Si in phengite.
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Omp þ Coe þ Rt is present in the field with an O content
of <0�50 mol %. A value of O¼ 0�40 mol % was chosen
for further modeling of both samples. For the kyanite-
bearing eclogite (WYKK-3), the modeled mineral as-
semblage of Grt þ Omp þ Ky þ Ph þ Coe þ Rt occurs
over the full range of O contents investigated at T above
�790�C (Supplementary Data Electronic Appendix Fig.
S2b). A low bulk-rock O content is suggested by the low
Fe3þ content of the rock-forming minerals in this sam-
ple (Cpx with Acm0–1 and Grt with Adr0–5). Thus, we use
a value of 0�30 mol % in further modeling of this sample,
which is similar to the other samples and consistent
with the mineral compositions.
Estimation of H2OTo estimate the appropriate H2O content for modeling,
we calculated T–H2O pseudosections at P¼ 4�0 GPa for
YK128-12, YKS-5 and WYKK-3. A T–H2O pseudosection
was not calculated for the bimineralic eclogite (YK5-2)
because all the rock-forming minerals (Grt þ Omp þCoe/Qz) are considered anhydrous in THERMOCALC.
For YK128-12, the range of H2O contents considered
is from 0�00 to 1�00 mol %. The calculated T–H2O dia-
gram is dominated by five-variant fields (Fig. 8). Free
H2O is present at T above �740�C and H2O above
0�26 mol %. Lawsonite-bearing phase assemblages de-
velop at T below �750�C and H2O above 0�26. The
observed mineral assemblage of Grt þ Omp þ Coe þPh þ Rt develops only in a narrow field with an H2O
content of �0�26 mol %. At H2O < 0�26 mol %, K-feldspar
develops because insufficient phengite is present to ac-
commodate all available K (phengite is the only hy-
drous mineral in the phase assemblage). Given the
likelihood that the peak mineral assemblage was fluid-
absent, the H2O content in this sample can be con-
strained only to be less than saturated; that is, less than
0�26 mol %. Although Kfs will be present in phase as-
semblage fields where H2O <0�26 mol %, the amount is
nominal (Kfs ¼ 0�27–0�00 mol % at H2O ¼ 0�20–0�26 mol
%) and its presence does not upset the topology of the
phase diagram; thus, Kfs may be neglected in the phaseassemblages (see Schorn & Diener, 2017). T–H2O pseu-
dosections calculated for YKS-5 and WYKK-3 give
Fig. 7. (a) T–O diagram calculated at P¼4�0 GPa for phengite-bearing eclogite YK128-12. (b) P–O diagram calculated at T¼800�Cfor YK128-12. The bulk-rock composition used is given in Table 5, with H2O in excess and O varying from 0 to 3 mol %. The fieldlabeled in bold type represents the observed peak assemblage of interest.
Fig. 8. T–H2O diagram calculated at P¼4�0 GPa for phengite-bearing eclogite YK128-12, using the bulk-rock compositiongiven in Table 5, with O¼0�25 mol % and H2O varying from 0to 1 mol %. The field labeled in bold type represents theobserved peak assemblage of interest.
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similar results, but different minimum H2O values
owing to different amounts of modal phengite
(Supplementary Data Electronic Appendix Fig. S3). For
YKS-5, the phase assemblage Grt þ Omp þ Coe þ Ph þRt þ Kfs develops in the phase assemblage field with
H2O of <0�96 mol %. For WYKK-3, the phase assem-
blage of Grt þ Omp þ Ky þ Coe þ Ph þ Rt þ Kfs devel-
ops in the phase assemblage field with H2O of
<1�45 mol %.
In summary, we find that the H2O content in the bulk-
rock compositions can be constrained only to a max-
imum value by phase equilibrium modeling to avoid
free water at the peak stage. This value corresponds to
all the water incorporated in the hydrous mineral phen-
gite. Any arbitrary value of H2O higher than this max-
imum will allow free H2O or additional hydrous
minerals in the modeled phase assemblage at the peak
stage, such as lawsonite, that are not present in the
samples of this study. However, any lower value will
promote the development of K-feldspar. Therefore, the
H2O contents chosen for further modeling are 0�26 mol
% for YK128-12, 0�96 mol % for YKS-5 and 1�45 mol %
for WYKK-3. Given the absence of hydrous minerals in
the peak assemblage of YK5-2, for further modeling the
H2O content was set to zero.
P–T diagrams with constrained O and H2Ocontents for eclogite-facies metamorphismPhengite-bearing eclogite YK128-12The P–T pseudosection for phengite-bearing eclogite
YK128-12 (Fig. 9a) was calculated in the NCKFMASHTO
system using the composition given in Table 5 (for
Fig. 9). Compared with Fig. 6, phase relations in the new
diagram are much simpler and are dominated by higher
variance assemblage fields (five- or six-variant). In
Fig. 9a, Law is absent because of the low H2O content
and kyanite is absent at P above �3�1 GPa owing to the
reduced O content in the bulk-rock composition.
Although Kfs is present in all fields, the modeled con-
tent is less than 0�01 mol % and may be neglected.
Plagioclase and Opx are present at low P (<2�0 GPa)
and high T (>660�C). The observed mineral assemblage
of Grt þ Omp þ Coe þ Ph þ Rt occurs at P above
�3�1 GPa, but over the full range of T from 600 to 900�C.
Grossular and XFe [¼ Fe/(Fe þMg)] isopleths for garnet
(0�29–0�33 and 0�66–0�72, respectively), Si isopleths for
phengite (3�43–3�58 p.f.u.) and j(o) isopleths for ompha-
cite (0�58–0�66) were calculated for the modeled P–T
range (Fig. 9b). In general, the Si in phengite, the gros-
sular in garnet and the j(o) of pyroxene all increase with
increasing P and record the highest values at the high-
est P shown. By contrast, XFe decreases with increasing
T, with isopleths that have negative slopes in kyanite-
bearing mineral assemblages at lower P but positive
slopes in kyanite-absent mineral assemblages at higher
P. In the phase assemblage field of Grt þ Omp þ Ph þCoe þ Rt þ Kfs, isopleths of grossular and j(o) are wide-
ly spaced and mineral proportions show little or no
change (less than 1 mol %) across the field, although
Fe–Mg exchange continues between garnet and
omphacite with increasing T.
The grossular content in a large garnet increases
from core to mantle from 0�29 to 0�33, and then
decreases from 0�33 to 0�30 at the inner rim. How much
these variations in composition have been flattened by
volume diffusion is unknown (Caddick et al., 2010).
Taking the changes in grossular content at face value,
they suggest an increase in P followed by a decrease.
Thus, the grossular content in the apparent core of the
garnet might record part of the prograde evolution.
Because primary kyanite occurs as inclusions in garnet,
it is likely that the prograde path passed through a
kyanite-bearing field.
Based on the petrography, the peak mineral assem-
blage corresponds to the field Grt þ Omp þ Ph þ Coe
(þ Kfs) in Fig. 9a. Although changes in the composition
of garnet and omphacite are small and variations in
mineral proportions are minor across this high-variance
field, restricting our ability to quantify the P–T condi-
tions more tightly (Fig. 9b), variation in the Si content in
phengite is dominantly due to changes in pressure.
Thus, we may use isopleths of Si content in phengite to
estimate a minimum P. In this sample, the highest Si
content of 3�54 p.f.u. was retrieved from phengite
included in garnet and from large phengite flakes. This
value corresponds to P of 3�7–4�7 GPa at T from 600 to
900�C (Fig. 9b), indicating that the peak P for this sam-
ple was >3�7 GPa at T of 600–900�C. The lowest Si con-
tent of 3�43–3�44 p.f.u. was retrieved from fine-grained
phengite, which corresponds to P of 3�8–4�2 GPa at
T> 660�C. We interpret the small decrease in P implied
by the change in phengite composition to record de-
compression during exhumation, consistent with the
decrease in grossular content of garnet from the mantle
(0�33) to inner rim (0�30) discussed above. XFe in garnet
varies from 0�67 to 0�70 from mantle to rim and the j(o)
in primary omphacite varies from 0�60 to 0�64. Isopleths
for these compositions define a field with P of 2�1–
3�4 GPa and T of 620–815�C (Fig. 9b). The mineral as-
semblage in this field is Grt þ Omp þ Ph þ Ky þ Coe/Qz
(þ Kfs), consistent with the observed assemblage in the
sample. Kyanite in this field was formed via the reaction
Grt þ Coe! Ky þ Omp during decompression.
Phengite eclogite YKS-5The P–T pseudosection for phengite eclogite YKS-5
(Fig. 10a) was calculated in the NCKFMASHTO system
using the composition given in Table 5 (for Fig. 10a).
Phase relations for this sample are similar to those for
YK128-12. The peak mineral assemblage field of Grt þOmp þ Coe þ Ph þ Rt (þ Kfs) is located at P above
�3�1 GPa at T of 600–900�C. In this field, the Si in phengite
occurring as inclusions in garnet and as coarse flakes out-
side garnet (3�47–3�58 p.f.u.) records P of 3�6–4�8 GPa at
T>640�C (Fig. 10a). Isopleths for XFe (0�59–0�61) and
grossular content (0�25–0�29) in garnet from core to
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mantle, as well as j(o) of omphacite (0�50–0�56), define a
field with P of 1�7–2�7 GPa and T of 670–760�C (Fig. 10a).
Kyanite-bearing eclogite WYKK-3The P–T pseudosection for the kyanite-bearing eclogite
(WYKK-3; Fig. 10b) was calculated in the NCKFMASHTO
system using the composition given in Table 5 (for
Fig. 10b). Figure 10b has been calculated up to P¼ 7�5 GPa
because the high XAl [Al2O3/(CaO þ NaO þ K2O) in mole
per cent] of the bulk-rock composition stabilizes kyanite
over a much larger range of pressures (up to 7�5 GPa over
the full T range) than in the phengite eclogites. The occur-
rence of kyanite as inclusions in garnet, as a rock-forming
mineral in the matrix and replacing garnet suggests that it
was stable during the prograde, peak and retrograde
stages. In Fig. 10b, isopleths for Si in phengite have rela-
tively steep positive slopes. In the field with the phase as-
semblage Grt þ Omp þ Ky þ Coe þ Rt (þ Kfs), the
isopleth for the highest Si in phengite (3�53 p.f.u.) occurs
at P of 3�9–5�9 GPa and T of 650–900�C (Fig. 10b). Isopleths
for XFe (0�64–0�66) and grossular content (0�26–0�32) in
garnet from core to mantle, and the j(o) of primary
omphacite inclusions in garnet and rock-forming ompha-
cite (0�51–0�59) define a field with P of 1�6–3�4 GPa and T
of 620–715�C in the mineral assemblage field Grt þ Omp
þ Kyþ Coe/Qz þ Rt (þ Kfs) (Fig. 10b).
Bimineralic eclogite YK5-2The P–T pseudosection for the bimineralic eclogite (YK5-
2; Fig. 10c) was calculated in the NCFMASTO system
using the composition given in Table 5 (for Fig. 10c). The
phase relations in this diagram are similar to those for
the phengite-bearing eclogites (Figs 9 and 10a), suggest-
ing that omission of K2O in the model system makes lit-
tle difference at the low concentration in this sample.
At P below 3�7–4�2 GPa, kyanite forms via the reaction
Grt þ Coe! Ky þ Omp during decompression, consistent
with the observed assemblage in this sample. The grossu-
lar content in garnet from an apparent core to the mantle
(0�31–0�33) corresponds to isopleths in the mineral assem-
blage field of Grtþ Ompþ Kyþ Coeþ Rt (Fig. 10c). Based
on petrological observations, the peak mineral assemblage
was Grt þ Omp þ Coe þ Rt, which is located in a field at P
above �3�7 GPa (Fig. 10c). In this field, neither mineral pro-
portions nor mineral compositions show much change.
The grossular content of garnet from the mantle to the
inner rim decreases from 0�33 to 0�29, corresponding to
decompression across isopleths in the Ky-bearing mineral
assemblage field (Fig. 10c). Given the garnet core–rim
structure shown by petrological observations, the grossu-
lar content from garnet core to mantle then to rim may re-
cord P–T information from the prograde, peak and
retrograde stages. Isopleths for XFe (0�67–0�71) and gros-
sular (0�29–0�33) in garnet from mantle to inner rim, and
the j(o) of primary omphacite (0�56–0�63) define a field with
P<3�7 GPa and T< 720�C in the mineral assemblage field
of Grtþ Ompþ Kyþ Coe/Qzþ Rt (Fig. 10c).
Assessing the presence of melt duringexhumationIn the literature, it has been suggested that low-volume
partial melting occurred during exhumation in the eclo-
gites at Yangkou Bay and at other localities in the Sulu
Belt (Zheng et al., 2011; Chen et al., 2014; Li et al., 2014;
Fig. 9. (a) P–T pseudosection calculated for phengite-bearing eclogite YK128-12, using the bulk-rock composition given in Table 5.(b) Shows calculated isopleths for grossular content (cg) and XFe (xg) in garnet, Si in phengite (Si) and j(o) of omphacite (jo). Thefields constrained by bold dashed lines indicate P–T constraints for near peak and retrograde stages, based on mineral assemb-lages and isopleths for the compositions of phengite and garnet and omphacite, respectively. The field labeled in bold type in (a)represents the observed peak assemblage of interest.
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Wang et al., 2014, 2016, 2017). In the veins in the
phengite-bearing eclogite YK128-12, we interpret sym-plectites of Kfs þ Bt þ Qz þ Pl around coarse phengite,
separating it from quartz, as evidence of partial melting
during decompression.
We use phase equilibrium modeling of this low-vol-
ume partial melting to provide an important constraint
on the low-pressure part of the P–T evolution by defin-
ing an appropriate solidus for crystallization of themelt generated during the late decompression stage.
To determine the solidus and assess the presence of
melt during exhumation, we calculated phase dia-
grams involving melt for YK128-12. The effective com-
position for modeling (Table 5) was calculated by
integrating an equal volume of coarse phengite and
quartz, using the molar volumes of phengite and
quartz from Holland & Powell (1998) and the averagecomposition of phengite in the veins from Table 4. The
pseudosections were calculated in the KFMASHTO
system. There is negligible sodium and no calcium in
the coarse phengite in the veins, so Na2O and CaOwere not included in the modeling. A–x relationships
(melt model for metapelite, biotite, orthopyroxene,
muscovite and cordierite) used in the modeling are
from White et al. (2014). The a–x model for melt is
based on the a–x model for haplogranite melt of
Holland & Powell (2001), which was calibrated up to
1�0 GPa. Accordingly, supra-solidus phase equilibriamodeled at P> 1�0 GPa should be interpreted with due
caution. Pure phases include K-feldspar, quartz, kyan-
ite, sillimanite and aqueous fluid (H2O).
The H2O content in the effective composition was
calculated from the average analysis of phengite in the
veins from Table 4. The presence of hematite, goethite
and barite in thin grain boundary veinlets in the eclo-
gites at Yangkou Bay demonstrates the presence of anoxidizing fluid during retrograde metamorphism (Wang
et al., 2016). In this study, the high Ae content in sodic
Fig. 10. P–T pseudosection calculated for phengite eclogite YKS-5 (a), kyanite-bearing eclogite WYKK-3 (b) and bimineralic eclogite YK5-2(c). The bulk-rock compositions are given in Table 5. Calculated isopleths for grossular content (cg) and XFe (xg) in garnet, Si in phengite(Si) and j(o) of omphacite (jo) are shown in the figures. The fields constrained by bold dashed lines indicate P–T constraints for near peakand retrograde stages, based on mineral assemblages and isopleths for the compositions of phengite and garnet and omphacite, respect-ively. The field labeled in bold type represents the observed peak assemblage of interest. (See online version for color images.)
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clinopyroxene against coarse quartz in the veinlets also
suggests that conditions during the retrograde meta-
morphism were relatively oxidizing. Therefore, the O
content in the effective composition was probably
higher than that used for the peak stage P–T pseudosec-
tion of Fig. 9. Following the method introduced above,
we calculated a T–O pseudosection at P¼ 1�2 GPa to de-
termine an appropriate O content (Supplementary Data
Electronic Appendix Fig. S4). The O content is con-
strained to be in the range of 0�0–0�4 mol %. Because
the O content was higher than that used for modeling
the peak eclogite-facies conditions (0�25 mol %), we
chose the maximum permissible value of O¼ 0�4 mol %
for modeling the P–T pseudosection discussed below.
The P–T pseudosection was calculated for P of 0�2–
1�6 GPa and T of 600–900�C (Fig. 11). Extrapolating the
a–x model for hydrous melt to very high pressures may
introduce errors (Holland & Powell, 2001), so we limit
the calculation to P of 1�6 GPa. In Fig. 11, the phase dia-
gram is dominated by five-variant fields. The solidus
occurs at T of 706–737�C, with the lowest T at
P¼ 1�47 GPa, where the slope changes from positive at
higher P to negative at lower P. At high P (>1�15 GPa)
and low T (<737�C), H2O is absent at sub-solidus condi-
tions. With increasing T at low to high P or decreasing P
at low to high T, Sill and then Ky replaces Ms via the re-
action Qz þ Ms ! Sill/Ky þ Bt þ Kfs þ H2O or Liq.
Cordierite is stable at T>640�C and P<0�7 GPa, and
orthopyroxene is stable at T> 814�C and P< 0�6 GPa.
In Fig. 11, calculated isopleths of mol % melt are
shown in the field of interest with the assemblage Bt þKfs þ Ms þ Qz þ Liq. The modeled melt content
increases with decreasing P (or increasing T) to a max-
imum of �8 mol % around 0�6 GPa and the isopleths be-
come slightly more closely spaced owing to the
progressive consumption of muscovite. We interpret
that the observed mineral assemblage of Kfs þ Bt þ Qz
[without Pl in the (Na2O þ CaO)-free system] that forms
the symplectites around the coarse phengite flakes in
the veins may have crystallized from melt under P–T
conditions corresponding to the solidus between the as-
semblage fields Bt þ Kfs þ Ms þ Qz þ Liq and Bt þ Kfs
þMs þ Qz þ H2O (Fig. 11). We estimate the symplectite
in the veins to be no more than 5% by volume. For a
modeled melt content of 1–5 mol %, P–T conditions at
the solidus are P of 0�9–1�3 GPa and T of 730–710�C
(Fig. 11).
P estimates using the average P–T methodIn addition to phase equilibrium calculations, the aver-
age P (avP) method of Powell & Holland (2008) was
used to estimate the P of equilibration for the main
rock-forming mineral assemblages in the eclogites at
Yangkou Bay. AvP was calculated using THERMOCALC
3.40 and the ds55 dataset (Holland & Powell, 1998;
updated November 2003). The equilibrated mineral
assemblages for calculation were Grt, Cpx, Ph, Rt and
Coe for YK128-12 and YKS-5, with Ky in addition for
WYKK-3. A–x models used were Grt (White et al., 2007),
Omp (Diener & Powell, 2012) and Ph (Coggon &
Holland, 2002). Rt, Coe and Ky are considered pure
phases, and fluid was assumed to be absent.
Calculations were performed at P of 1�5–5�5 GPa and T
of 600–900�C. Starting guesses for datafile coding of the
a–x relationships were calculated using the observed
mineral compositions, following the definitions for the
compositional variables for the minerals used in the a–x
relationships.As notified by E. Green (personal communication,
2017), when running THERMOCALC in AvP–T mode
there is a problem with the T-dependence of K.
Therefore, we calculated avP at a range of specified T, ra-
ther than AvP–T. A set of avP results using different start-
ing guesses shows that the maximum P can be obtained
using compositions of Grt with maximum mGrt zGrt2 (cor-
responding to maximum aPyr aGrs2), Cpx with maximum
jOmp/xOmp [jOmp corresponds to Na/(Na þ Ca) and xOmp
corresponds to Fe/(Fe þ Mg)] and Ph with minimum yPh
(maximum aCel), consistent with the logic outlined by
Carswell et al. (2000) for calculating the maximum
recorded pressure conditions for MT–UHP eclogite.
Results of the AvP calculations and the corresponding
standard deviations (SD) and sigfits are given in Table 6.
The pressures are similar for YK128-12 (4�106 0�44 GPa
at T¼800�C) and YKS-5 (4�15 6 0�47 GPa at T¼ 800�C),
but the pressure for WYKK-3 (4�406 0�5 GPa at
T¼ 800�C) may be slightly higher. AvP calculations for
Fig. 11. P–T pseudosection calculated for veins in the phengite-bearing eclogite YK128-12. The bulk-rock compositions usedare given in Table 5. Calculated isopleths for the amount ofmelt (green dotted line, in mol %) at supra-solidus conditionare modeled. The field constrained by a bold dashed line at thesolidus indicates P–T constraints for symplectite of Kfs þ Bt þQz þ Pl around coarse phengite in veins. The field labeled inbold type represents the observed peak assemblage of inter-est. (See online version for color images.)
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the mineral assemblage Di þ Hbl þ Pl þ Qz in the sym-
plectite in YK128-12 gave results with SD and sigfits too
high to be useful, which may imply disequilibrium.
DISCUSSION
The application of phase equilibrium modeling toUHP eclogitesO and H2OAs we have demonstrated in this study, it is important
to determine the appropriate O and H2O contents in the
bulk-rock composition as the first step in phase equilib-
rium modeling of UHP eclogites, as these variables in-
fluence the phase relations, and the proportions and
compositions of phases (Clarke et al., 2006; Rebay et al.,
2010). Although the Fe2O3 in the bulk-rock composition
can be determined analytically, studies have shown
that the oxidation state of eclogites, and, therefore, the
Fe2O3 content, may be affected by retrograde meta-
morphism (e.g. Li et al., 2016a). In addition, fluid cannot
simply be assumed to have been present in excess dur-
ing the metamorphic evolution, particularly for those
eclogites with evidence implying fluid-deficient condi-
tions, such as the survival of intergranular coesite and
the preservation of igneous mineralogy and texture in
the outcrops at Yangkou Bay.
For these reasons, it is necessary to use an iterative
process that combines T/P–O/H2O diagrams with the
observed mineral assemblages, and the mineral pro-
portions and compositions, accepting the caveats dis-
cussed above, to constrain O and H2O contents (Rebay
et al., 2010). Our phase equilibrium modeling shows
that the P–T stability of the common mineral assem-
blage Grt þ Omp þ Coe/Qz þ Rt 6 Ph 6 Ky, as well as
lawsonite-bearing assemblages, in MT–UHP eclogites is
dependent on the oxidation state and whether H2O is
present in excess. In this study, using reasonable values
for O and H2O in the phase equilibrium modeling, we
do not predict lawsonite to be stable at any stage of the
P–T evolution of the eclogites at Yangkou Bay, in agree-
ment with petrological observations. By contrast, based
on phase equilibrium modeling with H2O in excess, Wei
et al. (2013) argued that lawsonite was present at the
metamorphic peak in MT–UHP eclogites from the South
Dabie belt, despite its absence from the samples they
studied.
Isopleth thermobarometryThe advantage of pseudosection thermobarometry is
that it uses all the available information from the
samples studied—the bulk-rock composition, the phase
assemblage and, potentially, the phase proportions and
compositions (Powell & Holland, 2008). Thus, in prin-
ciple, the phase relations combined with compositional
isopleths for critical minerals may yield reliable P–T
constraints for different stages of the metamorphic
evolution.
For MT–UHP eclogites, the residence time at an ele-
vated temperature (>700�C) is likely to promote cation
diffusion and mineral re-equilibration (e.g. Chakraborty
& Ganguly, 1991; Caddick et al., 2010), potentially
resulting in elimination of information relating to the
prograde and/or peak stages of the metamorphism
from those rock-forming minerals that exhibit solid so-
lution, such as garnet and pyroxene. This problem has
been well characterized in granulite-facies metamorph-
ism and methods have been proposed to mitigate the
effects of diffusive re-equilibration (e.g. Frost & Chacko,
1989; Fitzsimons & Harley, 1994; Pattison & Begin,
1994; Pattison, 2003). The problem has not been as
widely appreciated in UHP metamorphism, but we note
that Baldwin et al. (2008) were not able to retrieve P–T
conditions in the coesite stability field using conven-
tional thermobarometry on eclogite with preserved
coesite. This failure to retrieve P–T conditions compat-
ible with the presence of coesite using conventional
thermobarometry was most probably due to post-peak
diffusive re-equilibration. Similarly, in pseudosection
thermobarometry, if the stability field for the peak min-
eral assemblage occurs over a wide range of P and T,
such as the Grt þ Omp þ Rt þ Coe 6 Ky 6 Ph (þ Kfs)
assemblages in this study, diffusive re-equilibration
may limit our ability to narrow down the range of P–T
conditions using isopleths of mineral composition, as
discussed by Wei et al. (2013).
Compositional isopleths commonly used in pseudo-
section thermobarometry of HP–UHP eclogites are the
divalent cations in garnet and Si in phengite, together
with the j(o) of omphacite (e.g. Wei et al., 2009, 2013;
Bruand et al., 2010; Groppo et al., 2015; Li et al., 2016b).
Although cation diffusion should be expected at
T> 700�C (Caddick et al., 2010), modified garnet growth
zoning can still be identified in MT–UHP eclogites from
both the Sulu and Dabie UHP belts (e.g. Enami &
Nagasaki, 1999; Schmid et al., 2000; Zhang et al., 2005),
including those with evidence of a granulite-facies over-
print (e.g. Chen et al., 1998), in eclogites from the
Variscan basement of the Bohemian Massif (O’Brien,
1997), and in migmatites from the Himalaya (Harris
et al., 2004) and the Variscan belt (Jones & Brown,
1990; Kalt et al., 2000). In these examples, the
Table 6: AvP results for the phengite-bearing eclogite YK128-12 and YKS-5 and the kyanite-bearing eclogite WYKK-3 at Yangkou
YK128-12 YKS-5 WYKK-3
T (�C) 700 750 800 850 900 700 750 800 850 900 700 750 800 850 900avP (GPa) 3�78 3�94 4�10 4�25 4�40 3�83 3�99 4�15 4�31 4�46 3�93 4�17 4�40 4�64 4�87SD 0�28 0�35 0�44 0�54 0�65 0�30 0�37 0�47 0�58 0�70 0�34 0�41 0�50 0�60 0�73Sigfit (rfit) 1�2 1�4 1�6 1�9 2�2 1�3 1�5 1�9 2�2 2�5 1�8 2�0 2�2 2�5 2�8
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preservation of modified growth zoning is interpreted
to be due to either the large grain size of the garnet or
decompression and fast cooling.
Other factors that may influence the retention of
compositional properties developed before or at the
metamorphic peak include diffusion kinetics of the ions
of interest, deformation, the presence of a fluid and the
oxygen fugacity (e.g. Chakraborty & Ganguly, 1991;
Caddick et al., 2010). Numerical modeling based on
experiments indicates that volume diffusion of Ca in
garnet is significantly more sluggish than that of the
other bivalent cations owing to its larger ionic size
(Chakraborty & Ganguly, 1991; Schwandt et al., 1996).
Therefore, Ca in garnet tends to better preserve pro-
grade chemical variation, whereas Mg, Fe and Mn are
commonly re-equilibrated and retain little or no chem-
ical zoning.
In this study, element mapping of garnet from two
of the samples shows internal zoning of Ca and, less
strongly, Mg and Fe, but not Mn (Fig. 4b and h). The
rate of exhumation of the eclogites at Yangkou Bay
and the fluid-deficient conditions may have
contributed to the preservation of Ca zoning in some
garnets. Although we are uncertain about the extent
to which the Ca zoning profiles in garnet have been
modified, the modeled Ca isopleths for garnet com-
positions from inner core to outer core from samples
YK128-12 and YK5-2, in combination with information
derived from mineral inclusions in garnet, suggest
preservation of a segment of the prograde path for
these eclogites (Figs 9b and 10c). By contrast, the
modeled Grs isopleths for garnet compositions from
mantle to inner rim in combination with the j(o) iso-
pleths for omphacite compositions indicate that re-
equilibration may have occurred along the retrograde
path (Figs 9b and 10a–c). Whether the zoning profiles
of these garnets have been modified during the sub-
sequent evolution of the eclogites cannot be
assessed, so the P–T results we present above based
on the zoning are taken as minima.
The Si content of phengite is more dependent on
pressure than temperature (e.g. Massonne & Schreyer,
1987; Schmidt et al., 1997, 2001). In experimental and
modeling studies of a variety of bulk compositions,
Fig. 12. P–T path for the eclogites at Yangkou Bay. The P–T path includes five stages based on phase diagrams for the four eclogitesmodeled, average P calculations and TiO2-in-phengite thermometry. Patches and the ellipse in different colors represent P–Tranges derived from different samples, corresponding to the P–T fields in the individual P–T pseudosections. The reaction Ph þ Cpxþ Coe/Qz þ (Ky)!Melt is from Liu et al. (2009b). (See online version for color images.)
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phengite exhibits systematic changes of Si content with
P (Hermann, 2003; Caddick & Thompson, 2008;
Hermann & Spandler, 2008; Auzanneau et al., 2010; Wei
et al., 2013). The increase in Si in phengite with increas-
ing P is due to a Tschermaks-type exchange where Mg
þ Si replaces Al between muscovite and celadonite in
mineral assemblages buffered by a SiO2 phase
(Hermann, 2003). However, during decompression,
change in the Si content of phengite is expected to be
limited, because change requires a coupled substitution
involving the tetrahedral site, which is rate limiting for
diffusion. Therefore, the Si in phengite is expected to be
a much more robust constraint on the highest pressure
preserved by the peak mineral assemblage than the Ca–
Fe–Mg proportions in garnet, where compositional
change during decompression involves only cation dif-
fusion on the dodecahedral site. Phase equilibrium
modeling of the samples in this study shows that at a
given T, Si in phengite depends mainly on P, whereas
changes in H2O and bulk-rock composition have little ef-
fect. Thus, the Si content of phengite is believed to be a
robust barometer that records reliable peak or near
peak pressures. To constrain the full P–T path for the
eclogites at the Yangkou locality, we rely on Si-in-
phengite isopleth barometry within the appropriate
phase assemblage fields for each of the three phengite-
bearing samples.
P–T path for UHP eclogites at Yangkou BayIf we accept the possibility that prograde compositions
may be preserved in the cores of some coarse-grained
garnets from the phengite-bearing (YK128-12) and bimi-
neralic (YK5-2) eclogites, we may define a segment of
the prograde P–T evolution using these data. Then by
combining the results from all four samples at UHP and
HP conditions for the retrograde P–T evolution, and
using the result from the symplectites around coarse
phengite flakes in the veins from sample YK128-12, we
may propose a full P–T path for the eclogites at
Yangkou Bay, from the prograde stage to the peak
metamorphic conditions and through several retro-
grade stages (Fig. 12).
The prograde stageBased on primary inclusions of Ky þ Omp þ Ph þ Coe
in garnet from YK128-12 and WYKK-3, the prograde P–T
path (Fig. 12, stage 1) passed through the assemblage
field Grt þ Omp þ Ky þ Ph þ Coe þ Rt (þ Kfs) (Figs 9b
and 10b). Although we cannot constrain the prograde
evolution within this assemblage field, increasing gros-
sular from inner core to outer core of garnet in YK128-
12 is consistent with increasing P and/or T through this
field. Accordingly, we propose a close to isobaric heat-
ing path at �2�8 GPa with a shallow dT/dP slope, fol-
lowed by a rapid compression to the peak metamorphic
conditions as shown in Fig. 12. Such a pattern is similar
to modeled P–T paths for the ocean crust in modern
subduction zones (Syracuse et al., 2010). In these
models, the low dT/dP paths follow a high dT/dP seg-
ment at intermediate depths, with of the order of a
300�C increase in temperature from 2�5 to 3�0 GPa. This
is attributed to heating at the plate interface at the tran-
sition from partial to full coupling between the slab and
the overlying mantle wedge in the subduction zone
(Syracuse et al., 2010). The similar change in the P–T
path during stage 1 is broadly consistent with prograde
Zr-in-rutile temperatures at 3�0 GPa of �680�C obtained
from rutile inclusions in garnet from eclogite at
Yangkou Bay (Wang et al., 2016).
The maximum P–T conditions retrievedThe peak metamorphic paragenesis in the phengite-
bearing (YK128-12 and YKS-5) and bimineralic (YK5-2)
eclogites comprises Grt þ Omp þ Coe þ Rt 6 Ph. In the
calculated P–T pseudosections these mineral assemb-
lages are stable over a wide range of P–T conditions at
P> 3�1 GPa and T of 600–900�C (Figs 9 and 10a, c), but
the proportions and compositions of garnet and
omphacite change little. By contrast, the Si content in
phengite changes systematically with P and, therefore,
may be used to estimate the P at or near the meta-
morphic peak (Fig. 12, stage 2).
In samples YK128-12, YKS-5 and WYKK-3, phengite
shows a limited range of covariation of Al and TiO2 con-
tents, and Mg/(Mg þ Fe) ratios with Si, implying some
resetting of phengite composition during decompres-
sion (Supplementary Data Electronic Appendix Fig. S1).
Nevertheless, a minimum value of the near peak P can
be estimated from the highest Si content in phengite in
each of the phengite-bearing eclogites. Phengite from
YK128-12 has a maximum Si of 3�54 p.f.u. and that from
YKS-5 has a maximum Si of 3�58 p.f.u., which give simi-
lar P of 4�0–4�6 GPa at T of 600–900�C, indicating that
the P should be no less than �4�5 GPa at T �850�C. By
contrast, the highest Si content (3�53 p.f.u.) in phengite
from the kyanite-bearing eclogite (WYKK-3) suggests P
of 3�9–5�9 GPa at T of 650–900�C (Fig. 10b). The Si in
phengite isopleths for each sample have different P–T
slopes. Thus, for an appropriate range of peak tempera-
tures of 800–900�C, we may estimate pressure based on
the crossing isopleths for the highest values of Si in
phengite, which yields a range from 4�2 to 5�9 GPa.
Given the possibility that the Si content in phengite may
have been reset during decompression, these pressures
probably represent minimum values, which allows the
possibility that the maximum P–T conditions may have
exceeded �5�5 GPa at >850�C.
Using avP, a similar range of P of 3�8–4�5 GPa at T of
700–900�C was retrieved from the assemblage Grt þ Ph
þ Omp þ Coe þ Rt in the phengite-bearing eclogites
YK128-12 and YKS-5, and a slightly higher P of 3�9–
4�9 GPa at T of 700–900�C was calculated for the mineral
assemblage Grt þ Ph þ Omp þ Ky þ Coe þ Rt in the
kyanite-bearing eclogite WYKK-3. The avP calculations
are broadly consistent with the pressures constrained
by the Si content in phengite. However, using isopleths
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of Grs and XFe in Grt with j(o) in Omp from the four
eclogites yields P–T results that are inconsistent with
those derived using Si in Ph (Figs 9b and 10a–c), which
may imply failure to maintain equilibrium between
these minerals during decompression. Therefore, the
avP results may not record the maximum P.
Disequilibrium plagues conventional thermobarome-
try owing to diffusional re-equilibration during retro-
gression, and the recovery of peak T is considered
unlikely (see Hacker, 2006). Although �Stıpska & Powell
(2005) used Grt–Cpx Fe–Mg exchange thermometry to
retrieve the peak T for an eclogite with the assemblage
Grt þ Omp þ Qz, the calculated uncertainty was of the
order of 100�C. Alternatively, experimental studies have
shown a strong positive correlation of TiO2 (or Ti) in
phengite with temperature when buffered with rutile
and quartz/coesite (Hermann & Spandler, 2008;
Auzanneau et al., 2010). Therefore, the TiO2 content in
phengite may be used to constrain the T at the peak or
near peak conditions for MT–UHP eclogite. The range of
TiO2 contents in phengite from all four samples in this
study suggests that the phengite compositions may
have been partially re-equilibrated during exhumation,
as also suggested by the range of Si contents in phen-
gite. The highest TiO2 contents in phengite are 0�71,
0�68 and 0�61 wt % for samples YK128-12, YKS-5 and
WYKK-3, respectively, constraining a maximum T of
800–900�C, at 4�5–5�9 GPa (Fig. 12).
In this study of the eclogites at Yankgou Bay, the max-
imum P–T conditions were potentially >5�5 GPa and
>850�C (Fig. 12, stage 2), which is higher than those esti-
mated by previous studies based on conventional thermo-
barometry, but closer to the P of >7�0 GPa at T> 1000�C
suggested by Ye et al. (2000), based on reintegration of
exsolved minerals with their host garnet. Furthermore,
our result is consistent within uncertainty with the min-
imum P of 5�5 GPa at T of 740–870�C retrieved by Liu et al.
(2006b) at Shanqingge, also in the Sulu Belt, but a few
hundred kilometers to the SW of Yangkou Bay.
The retrograde stagesEarly decompression (Fig. 12, stage 3) is recorded by
coarse phengite flakes and quartz in the veins from
YK128-12. The Si content of the coarse phengite ranges
from 3�53 to 3�44 p.f.u., corresponding to P of 4�5–
3�4 GPa (Hermann, 2003) or 4�0–3�2 GPa [equation (8) of
Caddick & Thompson, 2008] at T¼ 850–800�C (Fig. 12).
The TiO2 content in the phengite was not used to con-
strain temperature because rutile was absent and the
activity of TiO2 was unbuffered. Given the calculated
pressure, we speculate that quartz in these veins may
have been initially precipitated as coesite.
Subsequently, the UHP eclogites decompressed across
the coesite to quartz phase change (Fig. 12, stage 3 to
stage 4), facilitating re-equilibration of the main rock-
forming minerals.
During decompression and cooling (Fig. 12, stage 4),
isopleths for the Grs content of garnet, the XFe of the
garnet core or mantle and the j(o) of omphacite indicate
re-equilibration in the phase assemblage field Grt þOmp þ Ky þ Coe/Qz þ Kfs 6 Ph, consistent with the de-
velopment of secondary kyanite in YK128-12 (Fig. 2f),
WYKK-3 (Fig. 3e) and YK5-2 (Fig. 3i). These compos-
itional variations define a wide P–T range for each sam-ple (Figs 9b and 10), but whether the mineral
compositions represent equilibrium is uncertain and
our preferred P–T path passes through the high-T edge
of the range of P–T conditions determined.
Rock-forming phengite in the eclogites maintains a
euhedral shape and shows no evidence of instability,
such as breakdown by partial melting, implying that it
was stable throughout the post-peak P–T evolution. InFig. 12 we plot the experimentally determined solidus
for a phengite-bearing eclogite from the Dabie UHP ter-
rane (Liu et al., 2009b). This eclogite had a low H2O con-
tent (�0�29 wt %, essentially only the hydroxyl in
phengite and zoisite) such that initial melting of the
phengite-bearing assemblage occurred at relatively
high temperatures of 800–920�C, at P of 1�5–3�0 GPa.
The H2O content of the eclogites at Yangkou Bay is also
low (<0�5 wt %, Table 5), which could be the reason forthe stability of phengite during the decompression. This
expectation is consistent with the proposed P–T path,
which is located to the low-T side of the phengite-
breakdown melting reaction from Liu et al. (2009b).
Further decompression and local fluid migration
along veins and veinlets led to the development of a
symplectite of Kfs þ Bt þ Qz þ Pl around coarse phen-
gite in veins in YK128-12 and Hbl þ Pl 6 Qz 6 Cpx3 afteromphacite in veinlets in YK128-12 and YK5-2 (Fig. 12,
stage 5). Phase equilibrium modeling of the simplified
vein composition in YK128-12 yielded P–T conditions of
1�3–0�9 GPa at 710–730�C for the symplectite assem-
blage of Kfs þ Bt þ Qz (þ Pl) around phengite. These P–
T conditions provide the final constraint on the exhum-
ation P–T path for the eclogites.
During exhumation from HP eclogite- to amphibo-lite-facies conditions, the retrograde P–T path passed
through metamorphic conditions of P¼2�4–1�7 GPa at
T¼ 800–750�C and P¼ 1�3–0�9 at T¼ 730–710�C from
stage 4 to stage 5 (Fig. 12). The symplectites of Hbl þ Pl
6 Qz 6 Cpx3 after omphacite in veinlets and along the
edges of veins are interpreted to have formed by reac-
tion of local grain boundary fluid with the host under
amphibolite-facies metamorphic conditions, althoughthe P–T conditions cannot be constrained.
The final stage of the retrograde evolution is repre-
sented by coarse hornblende and plagioclase bordering
the veins in sample WYKK-3. Because the hornblende
and plagioclase compositions are different from those
in the symplectites along the veinlets in YK128-12 and
YK5-2, we infer that fluid migration along fractures
through the eclogite led to this retrogression at am-phibolite-facies conditions in lower crust.
In summary, this P–T evolution is similar to that in
many UHP terranes in the world (Ernst & Liou, 2008;
Liou et al., 2009).
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Uncertainties on the maximum P–T resultsIt is essential to assess the uncertainties when interpret-
ing the P–T results of phase equilibrium modeling
(including the avP calculations). Systematic and ran-
dom uncertainties should be propagated through all
calculations. Sources of systematic uncertainty include
the internally consistent dataset and mineral a–x rela-
tionships, whereas sources of random uncertainty are
mainly related to the analysis of minerals and determin-
ation of the effective composition of the bulk-rock or the
equilibration volume (Powell & Holland, 2008; Palin
et al., 2016). Normally both systematic and random
uncertainties need to be considered. However, when
comparing P–T results determined using the same
methods for several samples, systematic uncertainties
are common to all results, so that only relative errors
owing to random uncertainties need to be considered;
relative errors are generally small.
P–T conditions based on thermodynamic modeling
rely on rock, domain and mineral compositions. To min-
imize the uncertainties stemming from the electron
microprobe analysis of minerals, it is advisable to use
isopleths and independent sets of reactions that do not
involve Fe-bearing end-members (Hacker, 2006).
Uncertainties related to the effective composition may
be assessed using T/P–X pseudosections to investigate
those elements that influence the position of assem-
blage fields, and the mineral proportion and compos-
ition in phase diagrams (Powell & Holland, 2008).
For basic rocks, O and H2O are the main constituents
in the effective composition that affect the phase rela-
tions of high-variance assemblages at UHP conditions.
Using the appropriate O and H2O contents minimizes
the uncertainty associated with this composition. Thus,
in this study we calculated T/P–O/H2O pseudosections
to determine appropriate values for these two variables
and, based on the results of the modeling, we adjusted
the bulk-rock compositions accordingly.
For eclogite mineral assemblages such as Grt þOmp þ Coe/Qz þ Rt 6 Ky 6 Ph, the error derived from
mineral compositions is mainly due to uncertainty asso-
ciated with the ferric iron content in omphacite, which
might be 10% of the calculated result at UHP conditions,
whereas uncertainties associated with garnet and phen-
gite increase the compound error by only an additional
10% (Holland & Powell, 2008). In pseudosection ther-
mobarometry, the advantage of using isopleths of Si in
phengite, grossular in garnet and j(o) in omphacite is
that they involve no Fe-bearing end-members and the
uncertainty related to the Fe3þ in omphacite (and in gar-
net) becomes moot. In this study, the highest P was
retrieved from the Ky-bearing eclogite WYKK-3, where
Si in phengite in the assemblage Grt þ Omp þ Ky þCoe þ Ph (þ Kfs) indicates P> 5�5 GPa at T> 850�C. In
the analyses of the secondary standards, elements
were all within about 1% relative, with the exception of
Fe (4% relative). Therefore, the propagated uncertainty
on P constrained by Si-in-phengite barometry is
�0�20 GPa at T of 850–900�C.
For WYKK-3, the systematic errors on the P derived by
pseudosection thermobarometry are �0�07 GPa (at 1r) for
Si ¼ 3�53 p.f.u. in Ph, �0�07 GPa for j(o) ¼ 0�66 in Omp and
�0�12 GPa for z(g) ¼ 0�32 in Grt. However, we note that
these errors are based on propagation of uncertainties on
the thermodynamic data of the end-members of the
phases and do not include those associated with formula-
tion of the a–x relationships of the phases (see tutorials on
THERMOCALC website). The absolute errors on P derived
by combining the systematic and random uncertainties
discussed above for sample WYKK-3 are �0�3 GPa for the
maximum P of 5�5–5�9 GPa at T of 850–900�C; this esti-
mate should be taken as a minimum. Errors for the other
two phengite-bearing eclogites (samples YKS-5 and
YK128-12) are similar.
For avP calculations, the error on the results is
propagated from the uncertainties in the internally con-
sistent dataset, the interaction energies and the input
mineral compositions (Powell & Holland, 1994, 2008).
For WYKK-3, the maximum avP result of 4�4–4�6 GPa at
T of 800–850�C gives absolute errors of 0�5–0�6 GPa (at
1r; Table 6). The errors for the maximum avP calcula-
tions for the other two samples (YK128-12 and YKS-5)
are similar (Table 6).
In addition, differential re-equilibration of the main
rock-forming minerals during decompression may intro-
duce uncertainties into the results. Therefore, the errors
discussed above are almost certainly underestimated and
the results should be taken as minimum estimates.
Behavior of fluid during decompressionThe survival of intergranular coesite in eclogite and
preservation of relict igneous mineralogy and texture in
the protolith gabbro and surrounding gneisses at
Yangkou Bay (Liou & Zhang, 1996; Wallis et al., 1997;
Zhang & Liou, 1997; Wang et al., 2018), combined with
the conservation of pre-metamorphic O-isotope signa-
tures at General’s Hill (<2 km south of Yangkou; Wang
et al., 2017) indicate that the UHP metamorphism
occurred under fluid-deficient or fluid-absent conditions
(see Schorn & Diener, 2017). Indeed, the survival of
intergranular coesite requires essentially ‘dry’ condi-
tions (Mosenfelder et al., 2005; Wang et al., 2018).
However, the common development of veins and vein-
lets in the eclogites during the early and intermediate
stages of the retrograde evolution implies the presence
and local migration of some fluid. We interpret this ap-
parent contradiction to mean that during exhumation
the eclogite at Yangkou Bay became domainal, with low
a(H2O) host-rock domains where garnet and omphacite
have partially re-equilibrated during decompression
and intergranular coesite has survived, and high a(H2O)
domains represented by the veins and veinlets where
new mineral assemblages were formed, and if coesite
was present in these domains it was inverted to quartz.
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As discussed above, it is likely that the coarse phen-
gite flakes in the veins were precipitated under UHP
conditions (stage 3 in Fig. 12) close to the second critical
endpoints (P of 3�0–3�5 GPa at T of 750�C) in basalt/gran-
ite–H2O systems (Hermann et al., 2006; Hack et al.,
2007; Mibe et al., 2011). At UHP conditions and
T> 750�C, the nature of the fluid depends on the solute
content and it might be called supercritical fluid or hy-
drous melt according to preference at high solute con-
tent (Hermann et al., 2006, 2013; Wang et al., 2017).
Hereafter, we will simply use the term fluid. The source
of the fluid that precipitated the coarse phengite is un-
certain. As proposed by Wang et al. (2018), it may have
been exsolved from the nominally anhydrous minerals
that incorporate water as both structural hydroxyl and
molecular water in point defects within the crystal lat-
tice (Xia et al., 2005; Johnson, 2006; Chen et al., 2007,
2011; Zheng, 2009; Zhang et al., 2011). However, the
amount of fluid is likely to have been limited. The sur-
vival of intergranular coesite only millimeters away
from the veins and veinlets suggests the fluid was gen-
erated locally and that migration of this fluid was lim-
ited. Nevertheless, the higher a(H2O) in the veins may
have promoted partial melting at lower pressure during
the retrograde evolution, as recorded by the symplec-
tites of Kfs þ Pl þ Bt þ Qz around coarse phengite in the
veins (Fig. 2c and d).
In the simplified system used to model the symplec-
tites (Fig. 11), crystallization of melt at the solidus would
develop the phase assemblage Kfs þ Bt þ Qz þMs. The
presence of Pl in the symplectites in the veins indicates
that a small amount of fluid probably migrated into the
veins, introducing sodium (6 Ca) into the melt. The so-
dium (and calcium) could have been a product of the
transition from Na-rich Cpx in the host to Na-poor Cpx
in the veinlets.
Tectonic implicationsBased on conventional thermobarometry and/or Ti-in-
zircon and Zr-in-rutile thermometry, three patterns of
exhumation P–T path have been proposed for the Sulu
Belt: isothermal decompression followed by decreasing
P and T (e.g. Zhang & Liou, 1997; Banno et al., 2000;
Nakamura & Hirajima, 2000; Zhang et al., 2009), decom-
pression with significant cooling (e.g. Nakamura &
Hirajima, 2010; Li et al., 2014), and decompression with
heating (e.g. Wang et al., 1993; Zheng et al., 2011; Li
et al., 2014; Wang et al., 2016). At Yangkou Bay, previ-
ous studies of the eclogites have resulted in proposals
for all three types of exhumation P–T path (Zhang &
Liou, 1997; Liu et al., 2009a; Wang et al., 2014, 2016).
This variety of P–T paths may be partly due to the
large differences in calculated temperatures for the
peak or retrograde stages. One explanation for these
differences may be the application of different thermo-
barometric methods, combined with uncertainty over
the amount of Fe3þ in the ferromagnesian minerals.
Alternatively, different units in the Sulu Belt could have
experienced different exhumation processes. Eclogites
from Weihai and Rongcheng in the northeastern Sulu
Belt record evidence of a granulite-facies overprint with
the development of Opx þ Pl þ Qz (e.g. Wang et al.,
1993; Banno et al., 2000; Nakamura & Hirajima, 2000).
By contrast, at Yangkou Bay and Taohang, and for sam-
ples from the CCSD, symplectites of Cpx þ Pl þ Hbl þQz indicate an amphibolite-facies overprint during the
last stage of the exhumation (e.g. Zhang et al., 2006;
Zhu et al., 2007; Nakamura & Hirajima, 2010; this study).
The rare occurrence of blueschists in the southern sec-
tor of the Sulu Belt indicates lower P–T conditions than
further north. Therefore, a common P–T path during ex-
humation should not be expected for the HP/UHP rocks
in the Sulu Belt.
The new P–T path defined in this study is unlike
those proposed previously for Yangkou Bay and other
localities in the central and northern Sulu Belt. Here we
suggest that the prograde P–T path comprised a seg-
ment with a high dT/dP gradient, followed by a segment
with a low dT/dP gradient, similar to the P–T paths mod-
eled for the ocean crust during active subduction at the
present day (Syracuse et al., 2010). Based on peak min-
eral assemblages and mineral compositions, meta-
morphism at Yangkou Bay is confirmed as MT–UHP
type, with maximum P–T of >5�5 GPa at >850�C. These
P–T conditions cannot be retrieved using conventional
thermobarometry, although slightly higher P–T condi-
tions have been proposed based on interpretations of
textural and mineralogical features in eclogites (Ye
et al., 2000). In addition, similar P–T conditions to those
at Yangkou Bay have been retrieved from marble in
other units in the Sulu Belt (Liu et al., 2006b). The ex-
humation P–T path shows a continuous decrease in P
and T to high-pressure amphibolite-facies conditions.
The maximum P of >5�5 GPa indicates that different
units in the Sulu Belt may have been exhumed from dif-
ferent portions of the subduction system, as has been
discussed in other orogens such as the Alps, where dif-
ferent P–T paths are registered in adjacent segments of
the same tectonic unit (e.g. Bucher et al., 2005; Groppo
et al., 2009). To address this issue, additional quantita-
tive thermobarometry should be conducted on other
units in Sulu in the future. Based on the success of this
study, phase equilibrium modeling of eclogites at other
localities in the Sulu Belt may be useful in addressing
the subduction and exhumation processes that oper-
ated along the length of the belt.
CONCLUSIONS
1. We have shown that it is necessary to constrain the
O and H2O contents in the bulk-rock composition
prior to calculating a P–T pseudosection for MT–UHP
eclogite, otherwise the P–T pseudosection may be
unsuitable for the intended study. The results of this
study show that the P–T stability of the commonly
developed mineral assemblages Grt þ Omp þ Coe/
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Qz 6 Ph 6 Ky in MT–UHP eclogites are dependent
on both the O and H2O contents.
2. A new P–T path is proposed for the eclogites at
Yangkou Bay, extending from the prograde evolution
through the peak UHP conditions to the HP and am-
phibolite-facies retrograde conditions. The modeled
peak phase assemblage of Grt þ Omp þ Coe 6 Ky 6
Ph (þ Kfs) is in agreement with the petrological obser-
vations. Modeled isopleths for Si in phengite com-
bined with TiO2 contents in phengite yield minimum
P–T conditions of P> 5�5 GPa at T> 850�C. Modeled
isopleths for the compositions of garnet and ompha-
cite indicate that re-equilibration occurred in the phase
assemblage field of Grt þ Omp þ Ky þ Coe/Qz (þ Kfs).
Further decompression led to the development of Kfs
þ Bt þ Qz þ Pl around coarse phengite in veins and
symplectites of Hbl þ Pl 6 Qz 6 Cpx after omphacite
in veinlets. The retrograde P–T path passed through
metamorphic conditions of P¼ 4�0–3�4 GPa at T¼ 850–
800�C to 2�4–1�7 GPa at T¼ 800–750�C, before exhum-
ation to shallower crustal levels at P¼ 1�3–0�9 GPa and
T¼ 730–710�C.3. During exhumation, the eclogites at Yangkou Bay
became domainal, with fluid-deficient host-rock
domains where intergranular coesite was preserved
and fluid-present domains (veins and veinlets)
where new mineral assemblages developed. In
veins where a(H2O) was higher, decompression led to
partial melting of coarse phengite þ quartz in veins
and the crystallization of Kfs þ Bt þ Qz (þ Pl) around
coarse phengite at the solidus. In veinlets, migration
of fluid into the host along grain boundaries pro-
moted the limited development of symplectites com-
prising Hbl þ Pl 6 Qz 6 Cpx replacing omphacite.
4. This study demonstrates that phase equilibrium
modeling combined with hybrid thermobarometric
methods allows reconstruction of the P–T path for
MT–UHP eclogites at Yangkou Bay and the monitor-
ing of fluid behavior during decompression. The
new P–T path contrasts with those proposed previ-
ously for this locality and other localities in the Sulu
Belt. Additional studies using phase equilibrium
modeling are required to recognize spatial differen-
ces in P–T evolution in the Sulu Belt and to relate
these differences to subduction and exhumation
processes for different HP/UHP units within the belt.
ACKNOWLEDGEMENTS
Reviews by Eleanor Green, Johann Deiner and Gisella
Rebay, plus comments by Editor Joerg Hermann, sub-
stantially improved the original paper.
FUNDING
This study was supported by the National Natural Science
Foundation of China (Grants 41502043, 41272225 and
41572182), the Fundamental Research Funds for the
Central Universities, China University of Geosciences,
Wuhan (Grants CUG-G1323511572 and CUGL150815) and
China Postdoctoral Science Foundation (Grant
2016T90742), the MOST Special Fund (MSFGPMR02), and
the Opening fund (GPMR210703) from the State Key
Laboratory of Geological Processes and Mineral
Resources, China University of Geosciences Wuhan. B.X.
has been supported by the China Scholarship Council for
the postdoctoral program in the University of Maryland,
College Park.
SUPPLEMENTARY DATA
Supplementary data for this paper are available at
Journal of Petrology online.
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