JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 8
PAGES 1533^1552
2009
doi:10.1093/petrology/egp040
The Oxidation State of Metasomatized Mantle
Wedge: Insights from C^O^H-bearing Garnet
Peridotite
NADIA MALASPINA*, STEFANO POLI AND PATRIZIA FUMAGALLI
DIPARTIMENTO DI SCIENZE DELLA TERRA, UNIVERSITA’ DEGLI STUDI DI MILANO, 20133 MILANO, ITALY
RECEIVED NOVEMBER 12, 2008; ACCEPTED MAY 26, 2009
ADVANCE ACCESS PUBLICATION JULY 13, 2009
Oxygen fugacity (fO2) is an important parameter in determining the
relative stabilities of phase assemblages.Whereas a number of studies
have been devoted to determining the redox state of low-pressure
assemblages in the mantle system, the fO2 of the supra-subduction
mantle wedge is still poorly known. An essential input for fO2 estimates is the determination of the ferric^ferrous iron content of key
mantle minerals such as garnet, which can be measured using the
‘flank method’ technique on an electron microprobe. We selected samples of orogenic peridotites from the ultrahigh-pressure Sulu belt
(Eastern China) and from the Ulten Zone (Italian Alps) for
detailed case studies; these correspond to slices of metasomatized
mantle wedge sampled at different depths. They are characterized
by the assemblage phlogopite þ magnesite þamphibole in equilibrium with olivine, orthopyroxene and Fe3þ-bearing garnet. The
‘flank method’ measurements indicate that these pyrope-rich garnets
contain Fe3þ/Fe up to 012^014. For peridotite mineral assemblages fO2 can be evaluated from equilibria involving the Fe3þ garnet component skiagite (Fe2þ3Fe3þ2Si3O12) on the basis of
Fe3þ^Al substitution on the octahedral site, which is sensitive to the
garnet oxidation state. We modelled non-ideal mixing of Al and
Fe3þ on the octahedral site and non-ideal mixing on the dodecahedral
site, with a symmetric regular solution model for reciprocal solid solutions of Ca^Fe2þ^Mg^Al^Fe3þ-garnet. This allowed us to calculate garnet-peridotite fO2, given the presence of Fe3þ in garnet. Our
results indicate that the Sulu and Ulten peridotites record high
oxygen fugacities (FMQ to FMQ þ 2) compared with garnet peridotite xenoliths from the sub-cratonic mantle equilibrated at similar
pressures. The determination of the oxygen fugacity of these
hydrate^carbonate-bearing garnet peridotites allowed us to estimate
the speciation of C^O^H metasomatic fluids derived from the subducting slab, which are enriched in CO2. The fO2 evaluation of the
Oxygen fugacity (fO2) is conventionally used in petrology
as a variable describing the oxygen chemical potential
(mO2). It controls the valence state of redox-sensitive elements such as iron (Eugster & Wones, 1962) and plays an
important role in determining the relative stabilities of
phase assemblages. The geochemical cycle of many light
elements, including hydrogen and carbon, is strongly
affected by the mutual stability of reduced or oxidized
mineral and fluid species (Luth, 1989, 1999), often occurring as volumetrically minor compounds in peridotitic
rocks. In multiphase systems oxygen fugacity is classically
determined by the distribution of Fe3þ in mineral phases
and by the Fe2þ/Fe3þ equilibria. At present very little is
known about the relationship between mO2 and the phase
assemblage in complex mantle systems bearing alkali and
volatile components. This is mainly due to the very limited
number of systematic high-pressure studies at controlled
oxygen fugacity (Gudmundsson & Wood, 1995) and to the
technical complexity of determining the iron oxidation
state in experimental products, as well as in complex natural phase assemblages. Recent advances in spectroscopic
*Corresponding author. Telephone: þ39 02-50315613. Fax: þ39
02-50315597. E-mail: [email protected]
ß The Author 2009. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oxfordjournals.org
metasomatized mantle-wedge peridotites, representing the oxygen
chemical potential O2, provides a first step in unravelling the relationship between O2 and the metasomatic phase assemblage.
C^O^H fluids; mantle wedge; oxygen chemical
potential; oxygen fugacity; skiagite
KEY WORDS:
I N T RO D U C T I O N
JOURNAL OF PETROLOGY
VOLUME 50
and microprobe techniques for determining the redox state
of transition metals on a micrometer scale (e.g. microMo«ssbauer, McCammon, 1994; ‘flank method’, Ho«fer et al.,
1994) are opening new frontiers in the quantitative assessment of redox processes in mantle rocks.
For peridotite mineral assemblages fO2 can be evaluated
from several equilibria involving Fe3þ-garnet components,
where Fe3þ occurs in octahedral coordination (Luth et al.,
1990). For the olivine þ orthopyroxene þ Fe3þ-garnet
assemblage one of these reactions is represented by
2 Fe
2þ
3 Fe
3þ
2þ
¼ 4 Fe
2 Si3 O12
2 SiO4
ðskiagiteÞ
ðfayaliteÞ
ð1Þ
þ2 Fe2þ 2 Si2 O6 ðferrosiliteÞ þ O2
which involves the ferric garnet component ‘skiagite’
containing both Fe3þ and Fe2þ (Luth et al., 1990;
Gudmundsson & Wood, 1995). However, the applicability
of equilibrium (1) has been limited by the lack of thermochemical data for the Fe3þ-garnet component skiagite and
of an appropriate solid solution model for this phase.
Despite the number of studies that have been devoted to
determining the redox state of sub-cratonic upper mantle
(Wood et al., 1990; Ballhaus, 1993; Canil & O’Neill, 1996;
Simakov, 1998; McCammon et al., 2001; Woodland &
Koch, 2003; McCammon & Kopylova, 2004; Frost &
McCammon, 2008; Creighton et al., 2009), the fO2 of the
supra-subduction zone garnet-bearing mantle wedge is
still poorly investigated (Brandon & Draper, 1996;
Parkinson & Arculus, 1999; Peslier et al., 2002). Subduction
environments are complex systems, being sites where the
fluid phases metasomatize and refertilize the upper
mantle by transferring elements from the slab to the
mantle wedge. Occurrences of hydrous minerals coexisting
with carbonates and C polymorphs (e.g. phlogopite þ
magnesite þ graphite/diamond) in mantle-wedge peridotites (van Roermund et al., 2002; Carswell & van
Roermund, 2005) provide evidence that such fluids are
represented by C^O^H solutions, derived by devolatilization of the slab.
As case studies, we selected samples of orogenic peridotites from the ultrahigh-pressure (UHP) Sulu belt
(Eastern China) and from the Ulten Zone (Italian Alps),
which correspond to slices of metasomatized mantle
wedge sampled at different depths (Obata & Morten,
1987; Wang & Liou, 1991; Zhang et al., 1995, 2000, 2007;
Nimis & Morten, 2000; Rampone & Morten, 2001;
Scambelluri et al., 2006; Yang et al., 2007; Malaspina et al.,
2009). They are characterized by the occurrence of phlogopite þ magnesite (Sulu peridotite) and of amphibole
(Ulten peridotite) in equilibrium with olivine, orthopyroxene and Fe3þ-bearing garnet. We determined the Fe3þ content in garnets by electron probe microanalyses using the
‘flank method’ and applied equilibrium (1) integrating
a new solid solution model for the fO2 calculations.
NUMBER 8
AUGUST 2009
The determination of the oxygen fugacity of these metasomatized garnet peridotites will allow us to estimate the
speciation of C^O^H metasomatic fluids and to speculate
on the mechanism of transfer of C^O^H components
from the slab to the mantle wedge.
A N A LY T I C A L A N D
E X P E R I M E N TA L M E T H O D S
We measured the Fe3þ/Fe ratio of natural garnets by the
‘flank method’ on wavelength-dispersive spectra acquired
by electron microprobe (Ho«fer et al., 1994, 2000; Ho«fer &
Brey, 2007). This method exploits the low-energy emission
lines behaviour of iron, which is sensitive to the chemical
bonding, crystal structure, and coordination polyhedra in
isostructural minerals. In the FeL X-ray emission spectra
of Fe-bearing minerals, the La and Lb peaks, and Lb/La
intensity ratios shift with changes of the iron oxidation
state. This is illustrated in Fig. 1, where the emission spectrum of the Fe3þ-garnet endmember andradite is compared with that of the Fe2þ-garnet endmember
almandine. With the ‘flank method’ the Lb/La intensity
ratios are measured in correspondence to those Lb and
La wavelength positions where the differences between
the spectra of pure ferric and ferrous iron-bearing samples
are most pronounced (Ho«fer & Brey, 2007). As shown in
Fig. 1, for Fe-bearing garnets the two measuring positions
Lb and La for the flank method correspond to the minimum and the maximum of the spectrum resulting from
the difference between andradite and almandine, in black.
Measurement of Lb/La at these positions gives higher sensitivity in comparison to the conventional method of using
the Fe La and Fe Lb peak maxima or peak area ratios
(Ho«fer & Brey, 2007). We calibrated the ‘flank method’ on
a JEOL 8200 Superprobe at the Dipartimento di Scienze
della Terra (University of Milan). The spectrometer calibration was carried out by searching the peak for FeKa
(ninth order) on a metallic iron standard at 25 kV accelerating voltage and 80 nA beam current using a TAP crystal
and the integral pulse height analysis mode. The FeKa
(ninth order) peak position is shown in Fig. 1 and compared with the Lb and La flank wavelengths. With respect
to the Lb and La peaks, the shape of FeKa (ninth order)
is very sharp, allowing a more accurate spectrometer shift
correction. The spectrometer position was adjusted using
a routine automatically run by the program PeakFitH
(donated by H. Ho«fer & J. Boerder, JEOL Germany
GmbH). The flank method measurements of Fe3þ/Fe
were performed at 15 kV and 60 nA on TAP crystals, with
the 300 mm slit. Counting time was set at 300 s on both
Lb and La measurements to increase the counting statistics. The quantitative Fe3þ/Fe in garnets was determined
by applying the correction for self-absorption. Selfabsorption is an interaction process between X-rays and
1534
MALASPINA et al.
METASOMATISM OF MANTLE WEDGE
Fig. 1. FeL X-ray emission spectra of almandine and andradite collected at 15 kV and 120 nA, together with the spectrum of the difference
between andradite and almandine (in black). The flank method measurement positions (FeLa and FeLb) correspond to the maxima of the difference spectrum. The ninth-order FeKa peak search of metallic iron, collected at 25 kV and 80 nA using the integral pulse height analysis
mode, is used for the spectrometer calibration.
the sample volume that depends on both the Fe concentration and the path length of the FeL X-rays within the crystal. An increased path length results from a greater
penetration depth of the electrons. The measured Lb/La
ratio versus the known Fe2þ content of the standards gives
a linear relationship. The deviation of Lb/La of the garnet
with unknown Fe3þ from such a relation ( ratio) is a
measure of the Fe3þ content of the sample [see Ho«fer &
Brey (2007) for details]. The accuracy of the flank method
has been demonstrated in previous studies on mantle garnets, where an error of between 0.02 and 0.04
for Fe3þ/Fe has been documented in samples with
8^11wt % total Fe (Ho«fer & Brey, 2007).
We tested the flank method on natural and synthetic
garnet endmembers with fixed Fe3þ/Fe. Natural andradite from Val Malenco, Italian Central Alps (Fe3þ/Fe ¼
1020 0007; H. Ho«fer, personal communication), natural
almandine from Collobrie'res donated by A. B. Woodland
(Fe3þ/Fe ¼ 0031; Woodland et al., 1995) and natural spessartine^almandine solid solution (Fe3þ/Fe ¼ 0101 0015;
H. Ho«fer, personal communiction) were selected. In Fig. 2
the Lb/La intensity ratios measured in these garnets are
plotted versus the Fe (wt %). The results are compared
with the equilines of Fe3þ/Fe calculated by Ho«fer & Brey
(2007) from a multiple linear regression fit of flank
method measurements on different series of synthetic
garnets.
The flank method was also calibrated on almandine
containing Fe3þ/Fe ¼ 0023, synthesized at the laboratory
of experimental petrology of the Dipartimento di Scienze
dellaTerra, University of Milan, and on a skiagitic majorite
Fe2þ3Fe3þ2Si3O12 with Fe3þ/Fe ¼ 029 002 (Fig. 2).
Skiagitic majorite was synthesized at 10 GPa and 11008C
in a Walker-type multianvil apparatus at the
Dipartimento di Scienze della Terra (University of Milan)
using tungsten carbide cubes of 32 mm edge length and
8 mm truncation edge length. Assemblies were composed
of pyrophyllite gaskets, a prefabricated Cr2O3-doped
MgO octahedron with 14 mm edge length and a zirconia
sleeving containing a graphite resistance heater. An MgO
sleeve was placed between the capsule and the graphite.
The capsule was made from 23 mm (outer diameter)
gold tubing and sealed by arc welding. Temperature was
measured by an axial S-type (Pt^Pt90Rh10) thermocouple
inserted in contact with the capsule and was considered
accurate to 208C. The duration of the multianvil experiments was 7^10 h.
Starting materials were obtained from a quenched slag
prepared by melting stoichiometric amounts of Fe2O3 and
SiO2 at 15008C in a gas mixing vertical furnace at the
Dipartimento di Scienze della Terra (University of Milan).
The Fe2O3^SiO2 mixed powder was pressed into pellets
and fused onto 5 mm loops of Pt wire suspended in the hotspot of the furnace. The correct Fe3þ/Fe ratio (040) was
achieved by controlling the fO2 of the furnace atmosphere
using a CO^CO2 gas mix. The value of fO2 yielding the
desired Fe3þ/Fe was estimated from Kress & Carmichael
(1988) and from Roth et al. (1987, fig. 6902). The composition of the CO^CO2 gas mixture was set using a Tylan
FC^260 flowmeters-group controller. The operation of the
1535
JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 8
AUGUST 2009
Fig. 2. Lb/La intensity ratios (flank ratios) measured in natural and synthetic garnet standards, plotted vs total Fe (wt %), compared with the
equilines of Fe3þ/Fe calculated from equation (3) of Ho«fer & Brey (2007). Alm, almandine; Ski, skiagite; And, andradite; Spes-Alm, spessartine^almandine solid solution; Pyr, pyrope; Spes, spessartine; Gros, grossular.
controllers and the oxygen partial pressure were further
monitored by an extractive ZrO2 oxygen sensor. Samples
were drop-quenched into a beaker of water at the bottom
of the furnace producing glass and slag. The Fe3þ/Fe
ratio of the glass and slag was checked by titration wet
chemistry [following the procedure of Bezos & Humler
(2005)] and was found to be within 10% of the nominal
values.
Major element compositions of garnet standards and of
garnets from the Sulu and Ulten peridotites were analysed
by wavelength-dispersive spectrometry using a Jeol 8200
Super Probe at the Dipartimento di Scienze della Terra
(University of Milan). Acceleration voltage was set to
15 kV, beam current was 15 nA and natural silicates were
used as standards: Mg^Al^Si on pyrope, Fe on almandine,
Ca on grossular, Mn on rodonite, Cr on chromite, Ti on
Ti-ilmenite. A PhiRhoZ routine was used for matrix correction. Mineral analyses were always performed using
detailed back-scattered electron images to control the
microtextural site.
SAMPLES
We selected metasomatized garnet peridotites from the
UHP Sulu belt (China) and from the Ulten Zone (Italian
Alps) for this study. They both represent portions of the
supra-subduction zone mantle wedge tectonically sampled
from different depths by the continental crust during
subduction and/or exhumation (Brueckner, 1998;
Brueckner & Medaris, 2000). The major element compositions of the main phases and the ‘flank method’ Lb/La
ratios together with the Fe3þ/Fe contents of garnets are
reported in Table 1.
Magnesite^phlogopite-bearing garnet
peridotite (Sulu, Eastern China)
This mantle-derived garnet peridotite comes from
Donghai County, located at the southeastern end of the
Sulu UHP terrane. Lenses of garnet peridotite and pyroxenite are hosted by granitic gneiss and minor jadeitite,
quartzite and marble. Based on several petrological and
isotopic studies, it has been extensively demonstrated that
the Sulu UHP garnet peridotites correspond to slices of
supra-subduction zone mantle wedge tectonically
emplaced into the crust (Zhang et al., 1995, 2000; Yang &
Jahn, 2000; Yang, 2003). The analysed sample shows two
distinct mineral assemblages. The older one consists of porphyroclastic garnet (Grt1), coarse exsolved clinopyroxene
(Cpx1) and coarse phlogopite flakes (Phl1). The younger
paragenesis is shown in Fig. 3a and consists of finergrained olivine þ clinopyroxene (Cpx2) þ orthopyroxene
magnesite phlogopite (Phl2) equilibrated with neoblastic garnet (Grt2). Detailed petrological studies were
reported by Zhang et al. (2007) and Yang et al. (2007), who
indicated equilibration conditions of these peridotites at
pressures of 5^6 GPa and temperatures of 800^9508C on
1536
MALASPINA et al.
METASOMATISM OF MANTLE WEDGE
Table 1: Major element compositions (wt %) of the main phases and flank L/L ratios of garnets from the Sulu and
Ulten peridotites
Sulu
Grt1 core
Grt1 rim
Grt2 core
Grt2 rim
Cpx1y
Cpx2y
Oly
Opxy
SiO2
4204(018)
4245(008)
4219(030)
4230(021)
5498
5528
4060
5801
TiO2
001(001)
003(001)
001(001)
001(001)
003
016
b.d.l.
000
Al2O3
2284(005)
2268(009)
2261(049)
2196(041)
195
101
001
023
Cr2O3
227(003)
262(005)
178(064)
195(145)
182
201
001
003
146(065)
Fe2O3
056(037)
048(020)
082(025)
FeO
837(034)
960(010)
976(032)
964(050)
MgO
1914(026)
1856(003)
1854(012)
1804(032)
MnO
040(003)
048(001)
057(005)
066(006)
000
CaO
479(002)
496(005)
472(009)
468(020)
2160
NaO
n.a.
n.a.
n.a.
n.a.
234
10058
Total
Lb/La
10042
086(001)
10186
089(000)
10100
10070
091(002)
205
1576
242
1590
795
504
5011
3615
015
004
012
2134
b.d.l.
008
167
b.d.l.
b.d.l.
10047
9915
9967
091(001)
Fe
694(027)
779(010)
814(013)
852(009)
Fe3þ/Fe
007(004)
004(002)
007(002)
012(005)
Si
2989(0011)
2995(0007)
3000(0021)
3003(0010)
1984
2001
Ti
0000(0000)
0001(0001)
0001(0001)
0000(0000)
0001
0004
Al
1913(0006)
1885(0005)
1895(0033)
1850(0040)
0083
0043
0000
0009
Cr
0127(0002)
0146(0003)
0100(0036)
0101(0079)
0052
0058
0000
0001
Fe3þ
0030(0019)
0025(0011)
0044(0014)
0079(0035)
0000
0006
0000
0000
Fe2þ
0498(0021)
0566(0007)
0580(0019)
0576(0028)
0060
0064
0160
0140
Mg
2028(0024)
1952(0005)
1966(0014)
1921(0025)
0848
0858
1832
1850
Mn
0024(0002)
0029(0001)
0034(0003)
0040(0004)
0000
0005
0001
Ca
0365(0001)
0375(0003)
0360(0006)
0358(0017)
0835
0828
b.d.l.
0163
0117
b.d.l.
Na
n.a.
n.a.
n.a.
n.a.
Xpyr
068
067
067
066
Xalm
016
015
015
015
Xgro
013
013
013
013
Xski
003
005
005
006
0996
b.d.l.
1992
0000
0004
0003
b.d.l.
(continued)
the basis of several geothermobarometers involving pyroxenes and garnet.
Grt1 cores are characterized by pseudosecondary inclusions. A previous study reports the trace element compositions of these inclusions, showing enrichments in large ion
lithophile elements (LILE) with positive anomalies in Cs,
Ba, Pb relative to Rb and K, and high U/Th ratios
(Malaspina et al., 2009). This trace element signature is
related to the influx of a slab-derived incompatible element
and silicate-rich fluid, which metasomatized the garnet
peridotite during Triassic UHP metamorphism. The polyphase inclusions are remnants of the metasomatic fluid.
Similar enrichments are also recorded by the second generation paragenesis, pointing to a genetic relation between
the slab-derived fluids and Phl2 þ magnesite in equilibrium with neoblastic Grt2 and Cpx2 (Malaspina et al.,
2009).
As shown inTable 1, the older Grt1 is zoned with decreasing MgO and increasing FeO from core to rim. The neoblastic Grt2 is characterized by a major element
composition very similar to that of the Grt1 rims in terms
of MgO, CaO and FeO contents. Al2O3 decreases slightly
towards the rim, whereas Fe2O3 shows significant zonation
with variations from 082 to 146 wt %. Both garnets are
characterized by relatively high Cr2O3 concentrations,
reaching 34 wt % in Grt2 rims. Coarse Cpx1 and finegrained Cpx2 are diopsidic in composition. Cpx1 has
slightly higher Al2O3, Cr2O3 and Na2O than Cpx2.
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JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 8
AUGUST 2009
Table 1: Continued
Ulten
Ult12
Grt core
MK5C
Grt rim
Olz
Opxz
Grt core
Grt rim
Cpxz
Olz
Opxz
5618
SiO2
4181(024)
4135(005)
4084
5756
4210(044)
4180(037)
5423
4067
TiO2
006(002)
007(005)
n.a.
008
002(001)
002(001)
007
n.a.
003
Al2O3
2289(025)
2265(018)
n.a.
150
2282(010)
2283(011)
112
n.a.
162
Cr2O3
14(012)
140(002)
n.a.
017
029
n.a.
020
Fe2O3
118(016)
116(043)
FeO
920(031)
951(031)
162
919
609
MgO
1803(015)
1839(003)
4916
MnO
055(003)
055(000)
012
CaO
536(023)
517(001)
NaO
Total
10048
10024
128(005)
131(004)
144(020)
140(062)
878(018)
882(048)
3451
1809(015)
1815(015)
1814
4937
3500
012
062(003)
062(004)
007
014
012
n.a.
022
574(004)
574(006)
2405
001
021
n.a.
000
019
n.a.
n.a.
10016
10059
9978
9961
9945
979
641
10089
10070
Lb/La
092(000)
091(001)
090(001)
090(002)
Fe
812(004)
815(006)
783(009)
781(008)
Fe3þ/Fe
010(001)
014(004)
013(002)
014(005)
Si
2987(0005)
2968(0018)
0999
1971
2995(0010)
2982(0018)
1974
0998
1947
Ti
0003(0001)
0004(0002)
n.a.
0002
0001(0001)
0001(0001)
0002
n.a.
0001
Al
1927(0013)
1916(0005)
n.a.
0061
1913(0001)
1919(0008)
0048
n.a.
0066
Cr
0080(0018)
0079(0001)
n.a.
0005
0072(0003)
0074(0002)
0006
n.a.
0005
Fe3þ
0064(0009)
0063(0023)
0077(0011)
0075(0033)
Fe2þ
0550(0018)
0571(0015)
0200
0184
0522(0008)
0526(0027)
0054
0189
0176
Mg
1920(0021)
1967(0010)
1793
1761
1918(0010)
1930(0014)
0971
1806
1808
Mn
0034(0002)
0034(0002)
0002
0003
0038(0002)
0037(0003)
0002
0003
0003
Ca
0410(0015)
0397(0003)
n.a.
0008
0437(0004)
0439(0005)
0935
0000
0008
n.a.
0000
0013
n.a.
0000
Na
n.a.
n.a.
n.a.
n.a.
Xpyr
066
067
065
065
Xalm
017
016
017
017
Xgro
014
013
016
015
Xski
003
004
004
003
Garnet is normalized on the basis of 12 oxygens. Clino- and orthopyroxenes are normalized on the basis of six oxygens.
Olivine is normalized on the basis of four oxygens. Grt, garnet; Cpx, clinopyroxene; Ol, olivine; Opx, orthopyroxene.
yMalaspina et al. (2009).
zObata & Morten (1987).
As total iron.
Olivine and orthopyroxene have 92 mol % of forsterite
and enstatite, respectively, in agreement with the mantle
origin of the peridotite.
Flank method measurements have been performed on
both porphyroclastic Grt1 and porphyroblastic Grt2
(Fig. 3a). The average Lb/La ratios, Fe (wt %) and the
resulting Fe3þ/Fe corrected for self-absorption are
reported in Table 1. In Fig. 4a the measured Lb/La ratios
are plotted against the total Fe (wt %) and compared
with the Fe3þ/Fe regression lines calculated from equation (3) of Ho«fer & Brey (2007). To a first approximation,
the analysed garnets contain ferric iron and show
Fe3þ/Fe between zero and 01 for Fe ¼ 7^8 wt %. The
Fe3þ/Fe ratios of Grt1 and Grt2, quantitatively determined after the self-absorption correction, are plotted in
Fig. 4b vs Mg-number [¼ Mg/(Mg þ Fe2þ)]. In this diagram the Fe3þ/Fe of garnets from mantle xenoliths
sampled from the Olmani and Lashaine cinder cones
(northern Tanzania), kimberlites of the Kaapvaal craton
(southern Africa), the Udachnaya kimberlite pipe
(Siberia), the Western Gneiss Region (Norway) and Beni
Bousera (Morocco) are plotted for comparison [data from
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Fig. 3. Photomicrographs (transmitted light) of representative minerals in the garnet peridotites from the UHP Sulu belt and the Ulten Zone,
and Fe3þ analysis points (circles connected by dashed lines): (a) cross-polarized light image of an equilibrium texture between olivine, finegrained garnet, orthopyroxene and euhedral magnesite in the Sulu garnet peridotite; (b) plane-polarized light image of fine-grained garnet^
amphibole peridotite from the Ulten Zone. Ol, olivine; Grt, garnet; Opx, orthopyroxene; Mgs, magnesite; Amp, amphibole.
Canil & O’Neill (1996)]. Overall the Fe3þ/Fe ratios of the
garnets from the Sulu peridotites fall within the range of
worldwide mantle xenoliths (i.e. from 000 to 012). Grt2
has the highest values among these garnets, showing a
zonation of Fe3þ/Fe from 007 to 018 in the rim of some
grains. The porphyroclastic Grt1 is characterized by concentrations of Ca and Fe3þ (p.f.u.) similar to those of
garnet peridotites from southern Africa (Fig. 4c). Also,
Grt2 has Ca ¼ 035^036 (p.f.u.), and Fe3þ varies from
004 to 011 in the rims. Similarly, the Fe3þ/(Fe3þ þ Al) of
Grt1 in the Sulu peridotites is comparable with the range
for xenoliths from both African and Siberian kimberlites,
whereas Grt2 shows higher ratios towards the rims
(Fig. 4d).
Amphibole^garnet peridotite (Ulten Zone,
Italian Eastern Alps)
The Ulten Zone peridotites are enclosed as lenses
in gneisses and migmatites of the high-grade
Variscan Austroalpine basement (Italian Eastern Alps).
The country gneisses are associated with mafic rocks and
both record a metamorphic evolution from eclogiteto granulite- to amphibolite-facies conditions. The
peridotites form metre-scale to decametre-scale lenses
with textures changing from coarse porphyroclastic
spinel
lherzolites
to
fine-grained
amphibole þ
garnet-bearing lherzolites (Obata & Morten, 1987).
Phlogopite, dolomite, and apatite have also been recognized in some samples of the fine-grained garnet peridotites (Morten & Trommsdorff, 2003; Sapienza et al., 2009).
Most of the Ulten peridotite bodies record the transformation of lithospheric spinel lherzolites into
garnet þ amphibole and amphibole peridotites during a
continuous P^T history (Nimis & Morten, 2000). This
overall process was accompanied by intense deformation
and hydration of the peridotite slices to form neoblastic
garnet þ amphibole-bearing assemblages (Obata &
Morten, 1987).
The studied samples (provided by M. Scambelluri) consist of mylonitic garnet peridotites showing evidence of
fluid-induced metasomatism. They display a strong foliation marked by syn-tectonic pargasitic amphibole in textural equilibrium with garnet, olivine, ortho- and
clinopyroxene. Olivine and orthopyroxene are recrystallized and form a mosaic-like texture (Fig. 3b). On the
basis of the mineral paragenesis, Nimis & Morten (2000)
calculated the peak equilibration condition at 3 GPa and
8508C from the maximum stability field of hornblende þ dolomite (dolomite-in reaction: Wyllie & Huang,
1975; and hornblende-out reaction: Green, 1973). Detailed
bulk-rock and mineral trace element analyses of these samples show strong enrichments in Cs, Ba, Pb and U
and moderate enrichment in Li, indicating addition of a
fluid-mediated crustal component to the mantle rocks
(Rampone & Morten, 2001; Scambelluri et al., 2006).
The major element composition of garnet in both samples is homogeneous (Table 1). Similar to garnet from the
Sulu peridotite they are pyrope rich (67^68 mol %) with
relatively high Cr2O3 contents (125^140 wt %). The also
have high Fe2O3 contents, up to 160^185 wt % (Table 1).
Olivine is forsteritic (Fo 89^90 mol %) and orthopyroxene
is enstatite. Compared with the Sulu peridotites, orthopyroxene in both Ulten samples contains some Al2O3
(150^162 wt %), consistent with the lower pressure equilibration conditions of the Ulten peridotites (3 GPa).
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Fig. 4. (a) Flank-method Lb/La ratio measured in garnets from the Sulu and Ulten peridotites plotted vs total Fe content (wt %). The results
are compared with flank measurements on synthetic skiagite (SkiNM). The inset gives a close-up of the data showing averages of Lb/La intensity ratios vs Fe compared with Fe3þ/Fe regression lines. (b) Variation of average Fe3þ/Fe ratio vs average Mg number [¼ Mg/
(Mg þ Fe2þ)] in the cores and rims (indicated with C and R in the end of the sample number) of both generations of garnets from the Sulu peridotite and in garnets from the Ulten peridotite samples (Ult12 and MK5C). For comparison, data for garnets from mantle xenoliths originating
from northern Tanzania, southern Africa, Siberia, Norway and Morocco are also plotted (data from Canil & O’Neill, 1996). (c) Average variations of Fe3þ (p.f.u.) vs Ca (p.f.u.); (d) Fe3þ/(Fe3þ þ Al) vs Fe2þ (p.f.u.), indicative of skiagite^almandine substitution, for garnets from the studied peridotites compared with the reference data of Canil & O’Neill (1996). In all the figures average data are reported together with error bars.
Clinopyroxene occurs only in sample MK5C. It is diopside
in composition and is characterized by low Al2O3 (lower
than the coexisting enstatite) and Na2O below 02 wt %.
Flank method Lb/La ratios and Fe were determined
for garnets from both samples. The results reported in
Table 1 indicate that Fe3þ/Fe is between zero and 01 for
Fe 8 wt % (Fig. 4b), with an error range of 005.
Compared with garnets in the Sulu peridotites and in the
xenolith series studied by Canil & O’Neill (1996), their
Fe3þ/Fe ratios corrected for self-absorption reach the
highest values in sample MK5C (Table 1, Fig. 4b). Both
samples have Fe3þ (p.f.u.) and Fe3þ/(Fe3þ þ Al) values
very similar to those of the Grt2 rims in the Sulu peridotites (Fig. 4c and d). Only Ca is slightly higher,
044 p.f.u. in the clinopyroxene-bearing sample (Table 1,
Fig. 4c).
T H E R M O D Y N A M I C P RO P E RT I E S
O F Fe 2 þ3 Fe 3 þ2 S i 3 O 1 2 S K I A G I T E
A N D S OL I D S OLU T ION MO D E L S
It is beyond the scope of this study to derive a new, comprehensive thermodynamic model for garnet solid solutions. However, a consistent set of thermodynamic data
for Fe3þ-bearing garnets is required for the application of
equilibrium (1). This set can be retrieved based on existing
experimental data on skiagite and ab initio models of
garnet structure.
The reference molar volume V8, the isothermal bulk
modulus (KT) and the pressure derivative of the bulk modulus (K0 ) were experimentally determined by Woodland
& O’Neill (1993, 1999) and were used as a basis for constructing the thermodynamic dataset for skiagite
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Table 2: Elastic, thermal properties, Cp function, third law
entropy and Gibbs’free energy of skiagite
V81,298
3
(cm /mol)
12135
KT
(K 1)
(GPa)
1574
67
a0 105
a1 109
a2
2336y
70286y
02950y
Cp ¼ a þ bT þ cT 2 þ d T 2 þ eT 1/2
S8
(J/mol K)
(J/mol K) (kJ/mol)
a 102 b 10
8704y
c 105
d 105
Fe2þ 3 Fe3þ 2 Si3 O12 ¼ Fe3 O4 þFe2 SiO4 þ2 SiO2 :
G8
e 103
14476y 40045y 43207y 86526y 369871y
ð2Þ
G8 has been adjusted to fit the P^T location of reaction (2),
using the data for magnetite, fayalite and coesite from the
database of Holland & Powell (1998, updated 2002).
Calculations were performed using the Perplex computer
package (Connolly, 1990). The resulting reaction curve is
illustrated in Fig. 5b.
The almandine^skiagite solid solution was modelled by
ideal and non-ideal mixing of Al and Fe3þ on the octahedral site adopting a symmetric regular solution model.
The experimental data relative to the equilibrium
a ¼ a0 þ a1T þ a2T 2
K’
magnetite, fayalite and coesite at high pressure following
the reaction
Fe2þ 3 Fe3þ 2 Si3 O12 þFeAl2 O4 ¼Fe3 Al2 Si3 O12 þFe3 O4 ð3Þ
407202z
Woodland & O’Neill (1999).
yOttonello et al. (1996).
zFitted to data of Woodland & O’Neill (1993).
(Table 2). The isobaric thermal expansion equation used is
from Ottonello et al. (1996). Up-to-date calorimetric experiments on skiagite have not been reported and the CP function needs to be evaluated. Ottonello et al. (1996)
presented a static lattice energy and vibrational energy calculation for garnet structures following Kieffer’s model.
Thermal properties including heat-capacity functions
were described for garnet end-members. As an alternative,
a polynomial for CP can be obtained by summing the heat
capacities of single oxide components in the proportions
in which they occur in skiagite, as formulated by Berman
& Brown (1985). A comparison of the two estimates is
reported in Fig. 5a, showing that the function proposed by
Ottonello et al. (1996) is consistent with the high-temperature constraint defined by the Dulong^Petit limit
(Fig. 5a), CP ¼ 3nR þ a2VT/b, where a and b are the thermal expansion and the isothermal compressibility, respectively. Because the thermal dataset of Ottonello et al.
(1996) is internally consistent and includes the third law
entropy, it is also selected as a basis for retrieving the reference standard state molar Gibbs’ free energy (G8)
from experimentally determined high-pressure phase
equilibria.
Woodland & O’Neill (1993) synthesized the end-member
skiagite garnet together with the complete solid solution
series along the almandine (Fe2þAl2Si3O12)^skiagite
(Fe2þ3Fe3þ2Si3O12) join, where Al and Fe3þ mix on the
octahedral sites. Their experiments demonstrated that the
extent of Fe3þ substitution with respect to Al is strongly
pressure dependent. At 11008C, the formation of the endmember Fe2þ3Fe3þ2Si3O12 bounds the lower pressure stability of skiagite at about 93 GPa. Skiagite forms from
(Woodland & O’Neill, 1993), reported in Fig. 6, were fitted
adjusting the symmetric interaction parameters for both
almandine^skiagite and hercynite^magnetite solutions as
described by Luth et al. (1990) and Woodland & O’Neill
(1993). Moving from an ideal mixing (grey continuous
line), the interaction parameters were optimized for the
range 05Xski503, where Xski ¼ Fe3þ/(Al þ Fe3þ), corresponding to the compositional range expected to be most
relevant for natural UHP garnets. Best fit is achieved at
Walm^ski ¼ ^3705 kJ and Wherc^mt ¼ 3705 kJ.
The application to peridotite systems implies addition of
Mg and, to a minor extent, Ca to the solution models.
The redox reaction (1) can be therefore expressed as
pyrope-rich garnet in equilibrium with forsterite-rich olivine and enstatite-rich orthopyroxene þ O2. A non-ideal
mixing on the dodecahedral garnet site was then also treated with a symmetric regular solution model, using the formulation given by White et al. (2001) for reciprocal solid
solutions of Ca^Fe2þ^Mg^Al^Fe3þ-garnet. Almandine,
pyrope, grossular and skiagite were selected as linearly
independent end-members. Andradite and koharite
(Mg3Fe2Si3O12) can be expressed as a combination of the
previous components: andradite ¼1 grossular þ1 skiagite
^ 1 almandine, and koharite ¼1 pyrope þ1 skiagite ^ 1
almandine. The interaction parameters are Walm^gr ¼15 kJ
and Wpy^gr ¼ 80 kJ (White et al., 2001), Walm^py ¼ 25 kJ
(Holland & Powell, 1998) and Walm^ski ¼ ^371kJ (this
work). Because of the lack of data on possible solid solutions between pyrope^skiagite and grossular^skiagite,
these have been approximated as ideal solutions, so that
Wpy^ski and Wgr^ski ¼ 0.
R E S U LT S A N D D I S C U S S I O N
Calculation of fO2 in garnet-bearing
assemblages and C^O^H fluid speciation
Oxygen fugacity can be calculated from several equilibria
between coexisting phases in mantle peridotites (Luth
et al., 1990). For garnet peridotites fO2 can be determined
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Fig. 5. (a) Comparison of the heat capacity (CP) function calculated by Ottonello et al. (1996) and of the polynomial function obtained using
the formulation of Berman & Brown (1985) for skiagite, compared with the Dulong^Petit law. (b) Formation reaction of skiagite, reproducing
the experimental P^Tconditions reported by Woodland & O’Neill (1996) at 11008C.
by equilibrium (1) (Gudmundsson & Wood, 1995;
Woodland & Peltonen, 1999). We performed calculations
using solid solution models for garnet (previous section),
olivine (forsterite^fayalite binary mixture, Holland &
Powell, 1998), orthopyroxene (enstatite^ferrosilite mixture,
Holland & Powell, 1996), and clinopyroxene (diopside and
hedenbergite, Holland & Powell, 1996). We assumed negligible Fe3þ and Al partitioning in orthopyroxene, given
the low Al2O3 content in orthopyroxene coexisting with
garnet and clinopyroxene at ultrahigh pressures. The
resulting pseudo-univariant redox equilibria are shown in
Fig. 7. Two isobaric T^log fO2 sections were computed at 3
and 5 GPa, the equilibration pressures for the Ulten and
Sulu peridotites, respectively. The dashed curves represent
the equilibria selected for those pseudocompounds
approaching the mineral composition of the investigated
samples, where olivine and orthopyroxene consist of
89 mol % of forsterite and enstatite respectively (Fo89,
En89), and garnet is composed of 70% pyrope and 10%
grossular with variable skiagite and almandine contents
as a function of the oxygen fugacity. Equilibrium (1) is
therefore expressed as
Py70 Gr10 Al20x Skix
¼ Py70 Gr10 Al20ðx1Þ Skiðx1Þ
þFo89 þ En89 þ O2 :
ð4Þ
The x subscript (20) reflects the Al^Fe3þ exchange in the
octahedral site of the garnet as a consequence of the Fe2þ
oxidation in almandine.
Based on our solution model, the Fe3þ measured in garnets from the Sulu and Ulten peridotites is equivalent
to the content of skiagite substituting for almandine.
The content of the skiagite component in garnets from the
Sulu peridotite is of the order of 4 (1) mol % (Table 1),
and in the Ulten peridotite 3 (1) mol %. Cr-bearing
garnet components (uvarovite and/or knorringite), where
Cr substitutes for Fe3þ and Al in the octahedral site, have
not been considered here. Garnets from both peridotites
contain small amounts of Cr2O3 (Table 1), which were not
included in the pseudocompound calculation, resulting in
a possible overestimation of the skiagite concentration
(05^1mol %). Nevertheless, this variation can be
neglected in the fO2 calculation, as it is in the range of
the Fe3þ error (Table 1, Fig. 4c and d).
Oxygen fugacity calculated using the equilibria shown
in Fig. 7 ranges from ^862 to ^744 log units in the Sulu
peridotite (Fig. 7a) and from ^1069 to ^894 log units in
the Ulten peridotites (Fig. 7b), at the respective P^Tequilibration conditions. In Fig. 7 equilibria involving hematite,
magnetite, fayalite, quartz, ferrosilite, wu«stite and iron at
various buffering conditions are also shown for comparison
(continuous lines). The fayalite^magnetite^quartz (FMQ)
buffer is shown by the dotted curve, because of the metastability of this phase assemblage at UHP. However, it can
be useful to refer to such an equilibrium, to compare the
results of our study with the oxygen fugacities of garnet
peridotites from different geological settings, as discussed
in the following sections. In both samples the log fO2
[¼ log fO2(sample) FMQ] ranges between zero and
þ2. In particular, the Sulu peridotite shows a variation in
fO2 from the first generation (porphyroclastic Grt1) to the
second generation of garnets (neoblastic Grt2). This is
related to the different Fe3þ/Fe contents, which increase
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METASOMATISM OF MANTLE WEDGE
Fig. 6. Almandine^skiagite solid solution, modelled by mixing of Al and Fe3þ on the octahedral site, is compared with the experimental results
of Woodland & O’Neill (1993) as a function of pressure at 11008C. Open and filled symbols are the phase relations between garnet, spinel, hematite, pyroxene and SiO2 as reported by Woodland & O’Neill (1993, fig. 4). The crossed squares are re-equilibration experiments, the closed
square represents the garnet þ fayalite þ spinel þ quartz equilibrium, and the open square is the calculated reaction involving the end-members almandine ¼ hercynite þ fayalite þ quartz. The grey curves represent ideal mixing, whereas the black curves model a non-ideal mixing
of Fe3þ/(Al þ Fe3þ). grt, garnet; sp, spinel; hem, hematite; fa, fayalite; qz, quartz.
in Grt2 (Table 1, Fig. 4b). The flank method measurements
of the Sulu Grt2 also reveal a zonation in Fe3þ/Fe from
core to rim. This implies a log fO2 variation from
06 06 in the porphyroclastic Grt1 and in the Grt2 core
to 11 04 in the rim of the garnet neoblasts. As shown in
Table 1 and Fig. 4b, garnets from the Ulten peridotites are
characterized by Fe3þ/Fe ratios in the range of the Grt2
rims of the Sulu peridotite. The calculated log fO2 at
the P^T equilbration conditions of 3 GPa and 8508C
varies from 03 02 to 13 07 for sample Ult12, reaching
the highest values of 20 in sample MK5C.
In Fig. 7a and b, the redox reactions where clinopyroxene is stable are also plotted (dotted curves on the right of
the diagram) as a function of the garnet oxidation state.
These represent equilibrium (4), where a solid solution
between 90% diopside and 10 mol % hedenbergite (He10)
is involved in the reaction
Fo89 þ He10 þ Py70 Gr10 Al20x Skix
¼ Py70 Gr10 Al20ðx1Þ Skiðx1Þ þ En89 þ O2 :
ð5Þ
Based on our calculations, at 5 GPa the pseudo-invariant
equilibrium (5) occurs only at T49008C (Fig. 7a) and the
difference in the calculated log fO2 using equilibrium (4)
or (5) at this pressure is of the order of 05 log units, a
negligible value, within the error indicated by the vertical
bars.
The fO2 recorded by the Sulu and Ulten peridotites has
important implications for the speciation of the coexisting
C^O^H fluid phase. The occurrence of hydrous minerals
in equilibrium with carbonates has been widely reported
in mantle-wedge peridotites. The garnet peridotite from
the Sulu belt studied here is only one of several examples
of slices of mantle-wedge peridotite where the coexistence
of hydrous minerals and carbonates (phlogopite and
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Fig. 7. Pseudo-univariant equilibria calculated at 5 GPa (a) and 3 GPa (b) plotted as a function of oxygen fugacity and temperature. The
dashed curves and the dotted curves on the right represent equilibrium (4) and (5) respectively, selected for pseudocompounds consisting of
89% forsterite, 89% enstatite and Fe^Mg^Ca garnet solid solution. Isopleths indicate the garnet compositions, represented by 70% pyrope,
10% grossular (Py70Gr10) and variable almandine and skiagite percentage (Al20^xSkix), as a function of fO2. Invariant points refer to the
stability of enstatite and diopside. Hematite^magnetite, fayalite^quartz^magnetite, fayalite^magnetite^ferrosilite, wu«stite^magnetite and
iron^wu«stite buffers are also plotted for comparison. hem, hematite; mt, magnetite; fa, fayalite; q, quartz; fs, ferrosilite; wu, wu«stite.
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METASOMATISM OF MANTLE WEDGE
magnesite, Fig. 3a) provides evidence for metasomatism by
fluid phases enriched in CO2 and incompatible elements
(Carswell & van Roermund, 2005; Zhang et al., 2007;
Scambelluri et al., 2008; Malaspina et al., 2009). Such
fluids probably correspond to complex C^O^H solutions,
derived by dehydration and decarbonation of the slab.
Recent findings of OH-bearing minerals coexisting with
carbonates and C polypmorphs, such as phlogopite þ
magnesite þ graphite/diamond within polyphase inclusions
in garnets from mantle-wedge peridotites (van Roermund
et al., 2002; Carswell & van Roermund, 2005), indicate
that the fluid speciation is closely related to the oxidation
state of the system.
Although neither diamond nor graphite have been
found in the studied peridotites, the C-saturated system
can be used as a proxy to estimate the fluid composition,
when the graphite saturation boundary approaches the
binary H2O^CO2 and H2O^CH4 joins (Holloway &
Reese, 1974; Connolly, 1995). This occurs at high pressure
and relatively low temperature; that is, at conditions typical for the formation of the UHP peridotites investigated
here (Connolly, 1995). At UHP conditions C-saturated
fluids will be either CH4-bearing or CO2-bearing as a
function of the oxygen fugacity imposed by the redox conditions of the system. In both studied peridotites the
oxygen fugacities calculated from equilibria (4) and (5)
are above the FMQ buffer, with log fO240, and the
coexisting C^O^H fluid is enriched in CO2 (Connolly,
1995).
The oxygen fugacity of metasomatized
mantle wedge
The interpretation of the oxygen fugacities retrieved from
mantle rocks has been a subject of debate in recent years.
Thermodynamic calculations based on equilibria
between olivine^orthopyroxene^spinel (O’Neill & Wall,
1987) and olivine^orthopyroxene^garnet (Luth et al., 1990;
Gudmundssonn & Wood, 1995) suggest a relatively oxidized mantle, with fO2 between FMQ and the wu«stite^
magnetite (WM) oxygen buffer. In contrast, Ulmer et al.
(1987) proposed more reduced fO2 values, close to the
iron^wu«stite (IW) buffer. Systematic fO2 calculations for
mantle xenoliths from different geological settings, both
in the spinel and garnet facies, indicate that the upper
mantle is zoned (Daniels & Gurney, 1991; Ballhaus, 1993;
Ballhaus & Frost, 1994; Woodland & Koch, 2003). In addition, there are a number of studies in the literature that
reveal lateral fO2 variations related to different tectonic
settings (e.g. Woodland & Koch, 2003; Frost &
McCammon, 2008). These are summarized in Fig. 8,
where the ranges of oxygen fugacities for spinel and
garnet peridotites from various tectonic settings (grey
lines) are plotted as a function of equilibration pressure.
Calculated fO2 varies from 1 to 3 log units below the
FMQ oxygen buffer in abyssal peridotite, representing the
oceanic mantle lithosphere (Bryndzia & Wood, 1990); peridotite massifs have log fO2 from þ1 to ^2, with samples
from Beni Bousera reaching FMQ ^ 4 (Woodland et al.,
1992, 2006). Spinel peridotites from the subcontinental
lithospheric mantle, not reported in Figure 8, have variable fO2 ranging from FMQ ^ 1 to FMQ þ 2, with significant heterogeneities (Canil et al., 1990; Wood et al., 1990;
Brandon & Draper, 1996; Parkinson & Arculus, 1999).
Garnet peridotites from sub-cratonic mantle record the
lowest fO2 with values below FMQ ^ 2 (Woodland &
Koch, 2003).
An important aspect of oxygen fugacity variations in the
mantle is related to the systematic decrease of fO2 with
depth (Canil & O’Neill, 1996; Woodland & Koch, 2003;
Rohrbach et al., 2007; Frost & McCammon, 2008). As
shown in Fig. 8, considering the average fO2 values for
the different settings, fO2 appears to decrease by almost
three orders of magnitude moving from mid-ocean ridge
basalts (MORB) to abyssal peridotites and peridotite massifs, down to continental xenoliths from Lesotho and
South Africa, for which log fO2 reaches FMQ ^ 2 and
FMQ 3, respectively, at 4^5 GPa. Such a decrease in
oxygen fugacity follows the trend of equilibrium (1), which
is represented by the black continuous curve in Fig. 8. The
explanation of this behaviour is twofold. First, according
to Wood et al. (1990), at constant garnet composition the
negative slope of equilibrium (1) with respect to FMQ
(Fig. 8) is due to the negative iV for fO2-dependent
garnet-bearing equilibria. We report in Fig. 9a equilibrium
(1) as a function of fO2 and pressure. The isopleths of
almandine and skiagite (dotted lines) represent the
increasing Fe3þ^Al substitution in the octahedral site with
increasing oxygen fugacity. The plotted equilibria have a
negative slope, indicating that at constant Xski fO2
decreases when pressure increases. It is important to note
that the FMQ buffer is positively correlated with fO2 and
P, magnifying the iFMQ of almandine^skiagite solid
solution equilibria, which will therefore decrease strongly
towards higher pressures. Second, the modal amount of
mineral phases fractionating Fe3þ (spinel, garnet, majorite) increases with increasing pressure (Gudmundsson &
Wood, 1990; Wood et al., 1990; Woodland & O’Neill, 1993;
Canil & O’Neill, 1996; Woodland & Koch, 2003;
Rohrbach et al. 2007). At fixed bulk Fe2O3/FeO, the activity
of the ferric iron components, magnetite and skiagite,
therefore tends to drop as a result of progressive dilution
yielding to progressive lowering of fO2 with depth
(Rohrbach et al., 2007).
A striking difference in oxygen fugacity is evidenced
when anorogenic samples plotted in Fig. 8 are compared
with garnet peridotites from the supra-subduction zone
mantle wedge (black lines). Overall, both the Sulu and
Ulten peridotites plot in the range of the highest fO2
values (FMQ to FMQ þ 2). These results are comparable
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Fig. 8. Ranges and average values of oxygen fugacity relative to FMQ (ilog fO2 ¼ log fO2sample ^ log fO2FMQ) for the Sulu and Ulten garnet
peridotites (black lines) plotted as a function of pressure (modified after Frost & McCammon, 2008). They are compared with selected examples
of spinel and garnet peridotites from various ectonic settings, equilibrated at similar T conditions: oceanic mantle lithosphere (Bryndzia &
Wood, 1990), peridotite massifs (Woodland et al., 1992, 2006), garnet peridotite xenoliths from sub-cratonic mantle (Woodland & Koch, 2003)
and spinel-facies peridotites from subduction zones (Brandon & Draper, 1996; Parkinson & Arculus, 1999; Peslier et al., 2002). Curve (1) is the
fO2 calculated from equilibrium (1) for end-member skiagite. The CCO oxygen buffer (C þ O2 ¼ CO2) and the C^H2O join (X(O) ¼ 1/3) separating CH4- and CO2-rich aqueous fluids, calculated at 9008C, are also plotted. [Note the intersection of CCO with equilibrium (1) at 52
GPa and ilog fO2 12.]
with oxygen fugacities calculated for spinel peridotite
xenoliths from subduction settings (Fig. 8), which fall
between FMQ ^ 1 and FMQ þ15 (Brandon & Draper,
1996; Parkinson & Arculus, 1999; Peslier et al., 2002). It is
worth noting that with respect to garnet peridotite xenoliths from the sub-cratonic mantle, equilibrated at similar
pressure conditions, the Sulu and Ulten peridotites record
much higher oxygen fugacities with differences of 3^4 log
units. Furthermore, in contrast to what might be expected,
garnets from the Sulu peridotites have similar oxidation
states to the garnets of the Ulten peridotites formed at
lower pressure. This indicates that they do not follow the
same trend of fO2 decreasing with pressure along equilibrium (1) as recorded by the sub-cratonic xenoliths. We calculated fO2 based on equilibrium (4) for a complex
system at 9008C; the resulting skiagite isopleths are plotted
in Fig. 9b as a function of pressure. Comparing Fig. 9a
and b, it is clear that the addition of Ca and Mg components changes the slope of equilibria involving garnet, olivine and orthopyroxene. From this plot, at constant Fe3þ
garnet composition, a pressure change does not have the
same effect on fO2 as observed in the simplified chemical
system Fe^Al^Si^O^H. Moreover, in agreement with
Frost (1991), Fig. 9b shows that the substitution of Mg for
Fe2þ in iron silicates, with a consequent increase of Mg/
(Mg þ Fe2þ), stabilizes them to higher oxygen fugacities.
As an example, from equilibrium (1) calculated at 5 GPa,
a garnet with 70% almandine is stable with fayalite and
ferrosilite at fO2 ¼ ^8. At the same pressure, a similar
almandine/skiagite ratio in a garnet with 70% pyrope
and 10% grossular is given by Al14Ski6, which is in equilibrium with Fo89 and En89 at fO2 ¼ ^7. Figure 9b therefore
shows that any generalization on the evolution of fO2
with depth cannot neglect the influence of chemical
1546
MALASPINA et al.
METASOMATISM OF MANTLE WEDGE
Fig. 9. Isothermal P^log fO2 sections showing pseudo-univariant equilibria calculated from equilibrium (1) (a) and equilibrium (4) (b) at
9008C. The dashed curves represent the isopleths of almandine and skiagite from the Fe3þ^Al substitution in the octahedral site. Equilibria plotted in (a) are negatively sloping, indicating an fO2 decrease with increasing pressure, at constant Xski. The shape of equilibrium (4) calculated
in the more complex chemical system Fe^Al^Ca^Mg^Si^O^H (b), is different and the curves are almost isobaric. The reference hematite^magnetite, fayalite^quartz^magnetite and magnetite^wu«stite buffers cross the plotted equilibria, being positively correlated with pressure. Such an
opposite correlation implies that the iFMQ of almandine^skiagite solid solution equilibria will be magnified with increasing pressure.
Abbreviations are the same as in Fig. 7.
components such as Ca and Mg, which greatly affect the
energetics of the equilibria used to determine the redox
properties of mantle rocks.
Although it is generally believed that the mantle wedge
above subduction zones is oxidized as a result of metasomatism by a slab-derived fluid phase, the process responsible for relative oxidation and the actual oxidizing
capacity of slab-derived metasomatic agents are still disputed. Oxidized components are proposed to be transferred to the overlying mantle wedge by melts and/or
fluids coming from the subducting slab. Silicate melts and
Fe-bearing hydrous melts are likely candidates to produce
oxidation, as Fe2O3 is preferentially fractionated into the
melt phase (Frost & Ballhaus, 1998; Mungall, 2002). Water
itself has long been considered as an oxidizing agent
because it dissociates into oxygen and hydrogen. Oxygen
would form ferric iron whereas hydrogen escapes the
system. This reaction has been proposed by Brandon &
Draper (1996) as a possible explanation for the oxidation
of the mantle wedge; however, Frost & Ballhaus (1998)
showed that H2O cannot be assumed as an efficient oxidizing agent because of the very low H2O dissociation constant. Also, it has been pointed out that hydrogen cannot
diffuse away from the reaction site unless it is consumed
by interaction with the surrounding rock, resulting in the
formation of reduced zones. A similar argument applies to
melts, silica-rich hydrous fluids or supercritical liquids,
derived from devolatilization or partial melting of the subducted oceanic crust (Hermann & Green, 2001; Schmidt
& Poli, 2003; Kessel et al., 2005; Auzanneau et al., 2006;
Hermann et al., 2007), which are expected to be stopped at
the slab^mantle interface by reaction with the overlying
peridotite (Scambelluri et al., 2006; Malaspina et al.,
2006). However, the ubiquity of Fe3þ-bearing oxides as
daughter minerals in fluid inclusions and brines from serpentinized peridotites (Scambelluri & Philippot, 2001;
Scambelluri et al., 2001), and from mantle orthopyroxenites
at UHP (Malaspina et al., 2006), may suggest that the
residual aqueous fluids or melts may be carriers of oxidized
components away from the reaction front. Therefore, if
net bulk oxidation can be demonstrated, the aqueous component should be regarded only as a medium for oxidation.
OP E N QU E ST ION: I S T H E
M A N T L E W E D G E ‘ OX I D I Z E D ’ ?
As shown in Fig. 8, detailed oxygen barometric studies
have been carried out on abyssal peridotites, continental
lithospheric mantle samples and peridotite xenoliths from
the sub-cratonic mantle. On the other hand, because of
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VOLUME 50
the relative scarcity of samples, only a limited number of
studies have been performed on peridotites from suprasubduction mantle wedges (Wood & Virgo, 1989; Brandon
& Draper, 1996; Johnson et al., 1996). Moreover, there are
no data for the oxygen fugacity of garnet peridotites from
the deepest portions of the mantle wedge, mainly because
of the lack of suitable samples. The garnet peridotites studied here are therefore important witnesses of the processes
occurring in the deep sub-arc mantle, where subductionrelated metasomatism leads to refertilization and partial
melting feeding the source region of arc lavas. Studies on
the fO2 of arc lavas (e.g. Christie et al., 1986; Carmichael,
1991) have been carried out based on the assumption that
the oxygen fugacities of these magmas, ranging from
FMQ to FMQ þ 6, directly reflect those of their mantle
source regions. Accordingly, it has been inferred that the
sub-arc mantle is more oxidized than the shallow mantle
(Parkinson & Arculus, 1999). This conclusion is consistent
with the accepted paradigm that the mantle wedge is oxidized and its oxidation state is related to metasomatism by
oxidizing fluids or melts originating from the subducting
slab. On the other hand, as pointed out by Lee et al.
(2005), many arc-related mantle xenoliths, often inferred
to represent the source of arc lavas, record fO2 values that
are not high enough to be the source of the arc lavas.
Therefore oxygen fugacities determined in mantle peridotites are not necessarily indicative of the arc magma
source region.
The interpretation of oxygen fugacities calculated from
inverse modelling of phase equilibria in mantle peridotites,
as well as in magmatic rocks, should take into account the
fact that the fugacity concept in this context is just a conventional representation of oxygen chemical potential
(mO2). The variation of mO2 in multi-component systems
is not a simple increasing monotonic function of the
number of moles of O2 in the system. Variations in mO2, at
constant P and T, are the result of varying proportions of
all the constituent chemical components, and therefore
phase assemblages. Frost (1991) has suggested that
Fe2þ/Fe3þ can be misleading if considered as the sole monitor of fO2. Systematic experimental studies on the relations between oxygen buffers and phase assemblages are
mostly restricted to low pressure (e.g. Eugster, 1959;
Eugster & Wones, 1962). Although the value of mO2 in peridotitic rocks is constrained by olivine, pyroxenes, and
spinel/garnet, the relationship between mO2 and the phase
assemblages stabilized by the metasomatic processes is
entirely unknown. From a thermodynamic point of view,
the addition of components such as K2O, H2O and CO2,
and the formation of phlogopite, amphibole and carbonates may lead to variations in mO2 that are completely
unrelated to net whole-rock oxidation or reduction. To
demonstrate this concept we calculated the variation of
mO2 in the system FeO^Fe2O3^SiO2^KAlSiO4 at 2 GPa,
NUMBER 8
AUGUST 2009
11008C and H2O-saturated conditions (Fig. 10). Let us
assume a K-free bulk composition X (Fig. 10) falling in
the three-phase field ferrosilite^magnetite^quartz and
add progressively an ‘oxidizing’ metasomatic component
introducing potassium, ferric iron, aluminium, and silica
(point Y in Fig. 10). The bulk composition will move from
the ferrosilite^magnetite^quartz^sanidine field to the ferrosilite^magnetite^fayalite^sanidine field, up to the magnetite^fayalite^sanidine^annite field, along the path
traced by the black arrow shown in Fig. 10. This net bulk
‘oxidation’ path corresponds to a reducing path in terms of
‘oxygen fugacity’. The inset of Fig. 10 shows the corresponding variation in mO2, which moves from ^4619 kJ in
the ferrosilite^magnetite^quartz^sanidine field (1), to
4703 kJ in the ferrosilite^magnetite^fayalite^sanidine
field (2), down to ^4833 kJ in the magnetite^fayalite^sanidine^annite field (3). It finally increases again when it
reaches the compositionY in the magnetite^hematite^sanidine field. The relationship between the intensive variable
mO2 and the number of moles of O2 is not straightforward
and counterintuitive results might occur.
The determination of whole-rock oxidation degree is a
demanding task in most metasomatized garnet peridotites
because of the large number of phases that may incorporate both ferric and ferrous iron, and that may show compositional zonation. The Fe3þ partitioning among the
peridotite mineral phases is often neglected. Canil &
O’Neill (1996) studied the distribution of Fe3þ in garnet
peridotite mantle xenoliths characterized by a continuous
increase of Fe3þ/Fe in garnet with temperature. They
demonstrated that the increase of Fe3þ in garnet with
increasing temperature does not depend on the wholerock Fe2O3 content, but is rather the consequence of the
redistribution of Fe3þ from clinopyroxene into the garnet.
The Fe3þ clinopyroxene^garnet partitioning could explain
the high Fe3þ/Fe in garnets from peridotites equilibrated
at high T and P. This implies again that the Fe3þ enrichment in garnet is not necessarily indicative of high wholerock oxygen contents or of the interaction with more oxidized metasomatic agents (Canil & O’Neill, 1996).
A further indication of complexities in redox processes
in the mantle comes from studies on the V/Sc systematics
in peridotites (Canil, 2002; Lee et al., 2005). These
studies on mantle xenoliths showed that there is a discrepancy between oxygen fugacity calculated by O2 thermobarometry, the ‘barometric fO2’, and fO2 inferred from
V/Sc data on arc xenoliths, the last suggesting a surprising
homogeneity in the redox state of the asthenospheric
mantle, including subduction zone environments.
CONC LUSIONS
The ‘flank method’ measurements of garnets from mantlewedge peridotites from the Sulu ultrahigh-pressure belt
and the Ulten Zone reveal Fe3þ/Fe ratios up to 012^014.
1548
MALASPINA et al.
METASOMATISM OF MANTLE WEDGE
Fig. 10. Composition diagram in the system FeO^Fe2O3^SiO2^KAlSiO4 at H2O-saturated conditions showing stable assemblages at 2 GPa
and 11008C. The black arrow traces the oxidizing path from composition X, in the three-phase field ferrosilite^magnetite^quartz, towards Y
when a metasomatic component introducing K2O and Fe2O3 is added. The inset shows fO2 and mO2 variations (continuous-line and dashedline curves, respectively) as a function of T at 2 GPa, showing a decrease in these variables when the bulk composition moves from X to enter
the ferrosilite^magnetite^fayalite^sanidine field and then the magnetite^fayalite^sanidine^annite field. The numbers of equilibria in the inset
refer to the field numbers in the compositional tetrahedron. Kals, kalsilite; San, sanidine; the other abbreviations are the same as in Fig. 7.
fO2 calculations, performed with an improved thermodynamic solution model for skiagite-bearing garnets, suggest
that the Sulu and Ulten peridotites record much higher
oxygen fugacities (FMQ to FMQ þ 2) than garnet peridotite xenoliths from the sub-cratonic mantle equilibrated at
similar pressure conditions. Such high fO2 values are
accompanied by the occurrence of phlogopite þ magnesite
and amphibole dolomite in the Sulu and Ulten peridotites, respectively. Estimates of the speciation of coexisting
C^O^H fluids, assuming C-saturation, reveals that metasomatism operated via aqueous fluids, relatively enriched
in CO2.
These data might suggest that metasomatism in the
mantle wedge is related to bulk oxidation. However, we
have demonstrated that the variation in fO2 in multi-component systems is not a simple increasing monotonic function of the oxygen content in the compositional space.
High mO2 (and fO2) can be attained by lowering the bulk
oxygen proportion in the system, because the chemical
potential of oxygen, and therefore its conventional
representation in fO2 space, exhibits a complex variation
as a function of the variable phase assemblages developed
in metasomatized peridotites. The determination of the
fO2 of metasomatized mantle-wedge peridotites therefore
represents only the first step in unravelling the relationships between mO2 and phase assemblages in multi-component mantle systems.
AC K N O W L E D G E M E N T S
We thank M. Scambelluri for providing us the peridotite
samples from the Ulten Zone studied in this work, and R.
Compagnoni and S. Xu for guiding the field trip in
Dabie^Sulu (China). The advice of H. Ho«fer for the flank
method microprobe calibration and for the helpful clarification in Fe3þ measurements has been much appreciated.
Discussion with G. Ottonello in the early stage of the
manuscript and critical reviews by C. Ballhaus and
S. Turner significantly improved the paper. The financial
support
by
the
Italian
MIUR-Cofin
PRIN
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JOURNAL OF PETROLOGY
VOLUME 50
2007NCN7EZ_T108002 to the project ‘C^O^H fluids,
hydrates, carbonates and crust^mantle mass transfer in
subduction zones’ is acknowledged.
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