JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 2
PAGES 291–314
2002
The Fluid-absent Partial Melting of a
Zoisite-bearing Quartz Eclogite from
1·0 to 3·2 GPa; Implications for Melting in
Thickened Continental Crust and for
Subduction-zone Processes
KJELL P. SKJERLIE1∗ AND ALBERTO E. PATIÑO DOUCE2
1
DEPARTMENT OF GEOLOGY, UNIVERSITY OF TROMSØ, 9037 TROMSØ, NORWAY
2
DEPARTMENT OF GEOLOGY, UNIVERSITY OF GEORGIA, ATHENS, GA 30602, USA
RECEIVED MARCH 25, 2000; REVISED TYPESCRIPT ACCEPTED AUGUST 16, 2001
Fluid-absent melting experiments on a zoisite- and phengite-bearing
eclogite (omphacite, garnet, quartz, kyanite, zoisite, phengite and
rutile) were performed to constrain the melting relations of these
hydrous phases in natural assemblages, as well as the melt and
mineral compositions produced by their breakdown. From 1·0 to
3·2 GPa the solidus slopes positively from 1·5 GPa at 850°C to
2·7 GPa at 1025°C, but bends back at higher pressures to 975°C
at 3·2 GPa. The melt fraction is always low and the melt
compositions always felsic and become increasingly so with increasing
pressure. The normative Ab–An–Or compositions of the initial
melts vary from tonalites at 1·0 GPa to tonalite–trondhjemites at
1·5 GPa, adamellites at 2·1 and 2·7 GPa, and to true granites
at 3·2 GPa. At pressures <>2·5 GPa zoisite and phengite break
down more or less simultaneously. At 3·2 GPa and 1000°C zoisite
is unreacted whereas phengite is absent, so that the first formed melt
at these conditions is granitic. Our experiments show that if
sufficiently high temperatures (of the order of 1000°C) are attained,
zoisite- and phengite-bearing eclogites can produce small fractions
of silicic melts of a wide range of compositions. These melts are
rich in water and, probably, in Sr and other incompatible elements,
so that they can act as metasomatic agents in the mantle wedge.
INTRODUCTION
felsic melt; metasomatism
Zoisite or clinozoisite is present in many high-pressure
eclogites, and many occurrences are known from the
Scandinavian Caledonides (e.g. Holsnøy: Austrheim &
Mørk, 1988; Jamtveit et al., 1990; Western gneiss region:
Griffin et al., 1985; Seve Nappe: Kullerud et al., 1990).
Indeed, most eclogites worldwide contain minor amounts
of zoisite or other hydrous phases, according to the
compilations of Smith (1988) and Carswell (1990). During
prograde eclogitization, zoisite forms by breakdown of
the anorthite component of plagioclase in the presence
of a hydrous fluid phase. The origin of these hydrous
fluids is controversial. One possibility is that they are
introduced from below during continental collision, for
example, if wet continental sedimentary rocks are deeply
subducted. During subduction of oceanic crust zoisite
forms by prograde metamorphism of hydrothermally
altered oceanic crust, which in its upper levels contains
low-temperature hydrous Ca-rich phases (e.g. Poli &
Schmidt, 1997). At the gabbro to dyke transition zone
the temperature and pressure are high enough to allow
formation of clinozoisite and epidote as a result of hydrothermal circulation (see Skjerlie & Furnes, 1996, and
references therein).
Because zoisite and epidote are hydrous phases that
have been experimentally shown to coexist with melt
(e.g. Naney, 1983; Thompson & Ellis, 1994; Schmidt &
∗Corresponding author. E-mail: [email protected].
 Oxford University Press 2002
KEY WORDS: zoisite; dehydration-melting; orogenic thickening; subduction;
JOURNAL OF PETROLOGY
VOLUME 43
Thompson, 1996) the epidote group minerals could in
principle be involved in fluid-absent dehydration-melting
reactions. Zoisite may thus be a potentially important
phase for melt generation at high pressures in thickened
continental crust and in subduction zones. Experimental
studies of zoisite stability, including its melting relations,
have been performed mostly in model systems and under
fluid-present conditions (both water-saturated and waterundersaturated conditions). In contrast, fluid-absent zoisite melting experiments on natural starting materials are
scarce, so that very little is known about the fluid-absent
phase relationships of this mineral in natural rocks at
elevated pressures and temperatures.
REVIEW OF PREVIOUS
EXPERIMENTAL WORK ON
EPIDOTE AND ZOISITE AND
APPLICATION TO NATURAL
ECLOGITES
Fluid-bearing experiments
NUMBER 2
FEBRUARY 2002
systems at water-saturated conditions from 2·2 to 7·7
GPa, to determine the stability of hydrous phases in
subducting oceanic crust, and to constrain reactions that
result in the release of H2O to the mantle wedge. Their
experiments showed a large stability field for zoisitebearing assemblages. They also demonstrated that zoisite
may occur as a stable phase under water-undersaturated
conditions in eclogites. Their study suggested an upper
pressure stability for zoisite at >3·2 GPa under watersaturated conditions. At higher pressures zoisite is replaced by lawsonite-bearing eclogites at ‘low’ temperatures (T < >700°C) and dry eclogites at higher
temperatures. Poli & Schmidt (1997) further showed
that amphibole has a non-temperature sensitive upper
pressure stability of >2·3 GPa. They thus argued that
zoisite may be the most important hydrous mineral in
the 2·2–3·2 GPa pressure range under water-saturated
conditions in basaltic to andesitic bulk compositions.
Fluid-absent melting experiments
Several sub-solidus and super-solidus experimental studies have been performed on synthetic and natural epidoteand zoisite-bearing assemblages under fluid-present,
water-saturated and water-undersaturated conditions.
Super-solidus experiments have demonstrated beyond
any doubt that epidote and zoisite are magmatic phases
at high pressures in both mafic and felsic systems (i.e.
Naney, 1983; Schmidt, 1993; Thompson & Ellis, 1994;
Schmidt & Thompson, 1996). Schmidt & Thompson
(1996) showed that magmatic epidote has a wide stability
field in the tonalite system at water-saturated conditions
and f O2 buffered at NNO (nickel–nickel oxide). According to their experimental results, epidote dehydration
intersects the H2O-saturated solidus at approximately
500 MPa and 660°C. At higher pressures epidote exists
as a magmatic phase and its upper temperature limit
increases with pressure until the plagioclase to garnet
breakdown reaction is intersected at >1·3 GPa. At 1·3
GPa, Schmidt & Thompson (1996) determined that the
epidote stability field extends from the water-saturated
solidus at >630°C to >790°C. In the presence of garnet,
above 1·4 GPa, the upper temperature stability limit for
epidote has a steep negative Clapeyron slope. Schmidt
& Thompson (1996) also performed experiments at more
oxidizing conditions [hematite–magnetite (HM) buffer]
and showed that the epidote stability field was enlarged
down to 300 MPa. Similar experiments on a granodiorite
located the epidote-out reaction at 100 MPa, but the
maximum thermal stability is about 50°C lower toward
higher pressures.
Poli & Schmidt (1997) determined the sub-solidus
phase relations in natural andesitic and synthetic basaltic
Boettcher (1970) performed experiments under fluidabsent and water-saturated conditions in the CASH
(CaO, Al2O3, SiO2, H2O) model system and showed that
zoisite goes through dehydration-melting at pressures
<>800 MPa. He also argued that the thermal stability
of zoisite is strongly reduced in the presence of H2O,
and that the dehydration-melting reactions have positive
dP/dT slopes to 3·5 GPa (the highest pressure investigated). Above 2·0 GPa, the CASH dehydrationmelting reaction proposed by Boettcher,
zoisite + quartz = anorthite + kyanite + liquid
takes place at temperatures higher than 1050°C.
Thompson & Ellis (1994) performed water-saturated
experiments on the CMASH model system (CaO, MgO,
Al2O3, SiO2, H2O), and showed that zoisite is stable to
temperatures above the water-saturated solidus at high
pressures. Although they did not perform any dehydration-melting experiments, they calculated that zoisite could be involved in dehydration-melting reactions
with amphibole and quartz to yield anorthite (low P) or
garnet (high P) in addition to clinopyroxene and waterundersaturated melt. Their analysis suggested that zoisite
undergoes dehydration-melting in the presence of amphibole and quartz from >1·0 to >2·5 GPa at temperatures from >780°C to >820°C in the CMASH
system. In the absence of zoisite, amphibole undergoes
dehydration-melting at considerably higher temperatures. These calculations suggest that the presence
of epidote may cause fluid-absent melting in some bulk
compositions to occur at temperatures close to the watersaturated solidus. Skjerlie & Johnston (1996) performed
fluid-absent melting experiments on a crustal rock that
292
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
contained the hydrous phases biotite (16 vol. %), amphibole (15 vol. %) and epidote (13 vol. %) in addition
to plagioclase and quartz. Their results strongly suggested
that the thermal stability for biotite is lowered in the
presence of epidote through a dehydration-melting reaction such as
biotite + epidote + quartz = amphibole + garnet +
alkali feldspar + melt.
At 1·0 GPa this reaction produces 5–10 vol. % melt at
850°C and amphibole dehydration-melting produces an
additional 25 vol. % melt from 875°C to 925°C. Evidence
for the involvement of epidote in the melting reaction
was the observation that more amphibole formed in
the melting reaction than could be explained if all the
anorthite component of plagioclase broke down, and that
the glasses were fairly rich in CaO. The experimental
results of Skjerlie & Johnston (1996) therefore support
the work of Thompson & Ellis (1994) in that the presence
of epidote may lower the dehydration-melting temperature to those of amphibole and biotite at high pressure.
In the study by Singh & Johannes (1996) of dehydration-melting of tonalitic rocks, zoisite formed inside
plagioclase crystals at 1·2 GPa in a Fe-free phlogopite
+ plagioclase + quartz assemblage. When using an
assemblage with biotite composition in the range annite50–70 ( f O2 close to Co–CoO buffer) epidote formed
above 0·8 GPa. Epidote also formed inside the plagioclase
crystals and these were surrounded by alkali feldspar.
Singh & Johannes (1996) concluded that the chemical
conditions inside and outside the plagioclase crystals were
different.
Melting in natural eclogites
The high-temperature phase relations of epidote–zoisite
under fluid-absent conditions may have important bearings on deep crustal processes and on subduction-zone
processes. Several eclogites in Norway record temperatures close to and above the water-saturated solidus
(see Fig. 8, below) and many of these eclogites contain
zoisite. In the western Gneiss Complex of Norway eclogites probably formed in response to subduction of the
continent Baltica below Laurentia [see Austrheim &
Mørk (1988) for a discussion and other references], and
maximum P–T conditions are as high as 800°C and 3·0
GPa (Fig. 8; e.g. Cuthbert, 1995, E. K. Ravna, personal
communication, 1999). The P–T estimates increase towards the coast, and along the coastline high-P migmatites
have been described (Cuthbert, 1995). Possible high-P
migmatites have also been described from the eclogites
north of Bergen (Andersen et al., 1991). The eclogites of
the Tromsø area also contain small amounts of felsic
material that may represent the crystallized products of
high-pressure melts. Melting in the eclogites of western
Norway has always been discussed in terms of anatexis
in the presence of a water-rich fluid phase (e.g. Jamtveit
et al., 1990), but, in the light of the experimental results
discussed above, H2O-undersaturated melting should not
be excluded. If high-P zoisite dehydration-melting occurs
in nature, it is likely to produce small amounts of melt
owing to the low H2O content of zoisite and the high
solubility of water in silicate melts at high pressures.
Small-scale zoisite dehydration-melting in the lower
levels of thickened continental crust is potentially important. It is a well-known fact that the rheological
strength of a molten rock is dramatically reduced at high
melt fractions when the rock changes from matrix- to
melt-supported (i.e. van der Molen & Paterson, 1979).
However, as shown by Stevenson (1989), any partially
melted rock undergoing deformation is texturally unstable
because of small-scale redistribution of melt relative to
solid. Further, in crust undergoing deformation, melt will
inevitably localize in veins and the strength of the crust
will consequently be significantly reduced because strain
will tend to localize where melt is present (e.g. Davidson
et al., 1994; Tommasi et al., 1994; Rushmer, 1995). Thus,
the generation of small amounts of melt that segregate into
veins could potentially help to initiate and/or accelerate
orogenic collapse. At pressures higher than the amphibole
stability field, zoisite is likely to be the most important
water-carrier in mafic bulk compositions and its melting
may be important in promoting or accelerating destabilization of overthickened crust with following collapse
and exhumation of deep-seated rocks. Zoisite melting
may also be envisioned to occur during exhumation of
deep-seated rocks, in particular if exhumation is characterized by heating.
Small-scale zoisite melting can also be envisaged to
occur in subduction zones. Because zoisite is the main
carrier of Sr in plagioclase-free rocks and also contains
much of the light rare-earth elements (LREE; e.g. Hickmott et al., 1992; Nagasaki & Enami, 1998), the melts
are likely to be strongly enriched in these elements and
can be added to the mantle wedge and the melts may
thus act as a metazomatizing agent. Because of the
discussion above, it is most important that we understand
the melting behaviour of zoisite-bearing high-pressure
rocks.
The purpose of this study is to determine under which
P–T conditions zoisite undergoes dehydration-melting in
eclogitic assemblages, and to study the compositions of
the experimentally produced melts and solid phases. Our
main goal is to determine if zoisite dehydration-melting
is to be expected under those P–T conditions that can
be reached during overthickening of continental crust
and the following exhumation, and during subduction of
oceanic crust.
293
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 2
FEBRUARY 2002
Table 1: Characterization of starting material (Verpenesset eclogite)
Bulk rocka
Modeb
Cpx
Garnet
Garnet
cores
rims
46
32c
Zoisite
Kyanite
Phengite
4
8
49·96
55·98
40·95
41·31
40·42
—
—
53·22
—
TiO2
0·76
0·01
0·08
0·03
0·09
—
—
0·43
—
Al2O3
19·97
Fe2O3d
2·00
FeOe
4·03
1·98
16·78
16·89
MgO
8·15
11·62
10·56
11·98
MnO
0·11
0·01
0·33
CaO
12·72
17·48
9·53
Na2O
2·57
4·57
K2O
0·14
Total
100·41
8·41
22·63
—
23·08
<1
32·42
—
—
28·21
—
—
—
—
—
—
1·36
—
—
0·71
0·07
—
—
4·03
0·32
0·04
—
—
7·41
24·56
—
—
0·02
0·03
0·02
—
—
0·61
—
—
—
1
Rutile
SiO2
—
8
Quartz
—
—
—
—
—
—
—
—
—
—
—
10·54
—
100·06
100·86
101·05
98·96
—
—
97·77
—
a
Analysed by X-ray fluorescence.
Modal composition is calculated by a combination of mass balance and point counting thin sections.
c
Mode for garnet is the sum of cores + rims.
d
Fe2O3 determined by titration. Fe3+ is present in zoisite and in garnet.
e
Total Fe calculated as FeO.
All values are in weight percent; —, no analyses. Probe analyses of minerals are averages of 5–10 different analyses.
b
CHARACTERIZATION OF THE
STARTING MATERIAL AND
EXPERIMENTAL PROCEDURES
To understand the melting behaviour of zoisite-bearing
high-pressure rocks, we have chosen as starting material
a non-retrograded zoisite-bearing eclogite (Table 1) from
Verpenesset in West Norway kindly provided by Professor
E. K. Ravna. This starting material was chosen because
it contains primary high-pressure zoisite in addition to
kyanite and quartz. It is also important that there are no
signs of retrogression in the sample so that no water is
tied up in secondary phases. The starting material contains minor amounts of phengite, and it experienced
Caledonian eclogite-facies metamorphism at about
700°C and 2·5 GPa (E. K. Ravna, personal communication, 1998), probably related to continental subduction of Baltica below Laurentia at >420 Ma (e.g.
Austrheim & Mørk, 1988). The Al2O3-rich nature of the
starting material (Table 1) suggests that it might represent
a plagioclase-rich layer in a layered mafic intrusion.
Experiments on the Verpenesset eclogite were performed in piston cylinder apparatus at the University of
Tromsø (1, 1·5, 1·8, 2·1 GPa) and the University of
Georgia (2·7 and 3·2 GPa). The rock was crushed and
loaded into 1·3 mm Au capsules that were welded shut
after drying for 24 h in a 110°C oven. The capsules were
enclosed in NaCl that acted as the pressure-transmitting
medium in NaCl–MgO–graphite cells. Experiments were
run for a considerable length of time (Table 2), to achieve
as much reaction as possible. The capsules were weighed
after each run and discarded if weight loss was detected.
Glass analyses of selected capsules with tears always
yielded Cl-bearing glasses, so that the absence of Cl from
capsules with similar post- and pre-run weights is a
reliable indicator that there was no mass exchange with
the pressure medium during the run. The oil pressure
was monitored by Heise gauges at Georgia and by a
Heise-type gauge at Tromsø, and was converted to sample
pressure by the ratio of ram to piston areas. Pressures
are assumed to be accurate to within 50 MPa. Temperature was measured with type C thermocouples
(W5Re/W26Re) relative to an external electronic icepoint (OMEGA MCJ) and controlled by Eurotherm 808
regulators. Successful runs were polished for scanning
electron microscope (Tromsø) and electron microprobe
(Georgia) studies. Glass analysis were performed with
the JEOL JXA 8600 superprobe at the Department of
Geology, University of Georgia.
Alkali loss during probing of hydrous silicate glasses
has been investigated in other experimental studies in
which pools of glass large enough to allow multiple
analyses with different analysed areas and counting times
were available (e.g. Patiño Douce & Johnston, 1991;
Patiño Douce & Harris, 1998). Such large glass pools
were not produced in any of the experimental products
reported here. Because of this, the results obtained in
those earlier studies were assumed to be valid for this
294
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
Table 2: Phase assemblages and experimental run conditions
Sample
T
P
Duration
(°C)
(GPa)
(h)
Phase assemblage
KALB30
900
1
361
Cpx
Opx
Pl
Gt
Gl
KALB17
950
1
385
Cpx
Opx
Pl
Gt
Gl
KALB31
1000
1
142
Cpx
Opx
Pl
Gt
Gl
KALB34
1050
1
169
Cpx
Opx
Pl
Gt
Gl
KALB23
875
1·5
501
Cpx
Zo
Qtz
Pl
Ky
Gt
Gl
KALB19
900
1·5
382
Cpx
Zo
Qtz
Pl
Ky
Gt
Gl
KALB13
950
1·5
332
Cpx
Zo
Qtz
Pl
Ky
Gt
Gl
KALB20
1000
1·5
191
Cpx
Qtz
Pl
Ky
Gt
Gl
KALB24
1050
1·5
95
Cpx
Qtz
Pl
Ky
Gt
Gl
KALB21
900
1·8
188
Cpx
Zo
Qtz
Pl
Ky
Gt
Gl
KALB32
950
1·8
409
Cpx
Zo
Qtz
Pl
Ky
Gt
Gl
KALB33
1050
1·8
114
Cpx
Qtz
Pl
Ky
Gt
Gl
KALB6
800
2·1
645
Cpx
Zo
Ph
Qtz
Ky
Gt
KALB5
850
2·1
452
Cpx
Zo
Ph
Qtz
Ky
Gt
KALB34
900
2·1
115
Cpx
Zo
Pl
Ksp
Qtz
Ky
KALB1
950
2·1
118
Cpx
Zo
Pl
Ksp
Qtz
Ky
Gt
KALB22
975
2·1
114
Cpx
Zo
Pl
Qtz
Ky
Gt
Gl
KALB7
1000
2·1
89
Cpx
Zo
Pl
Qtz
Ky
Gt
Gl
KALB9
1025
2·1
98
Cpx
Zo
Pl
Qtz
Ky
Gt
Gl
KALB10
1050
2·1
78
Cpx
Zo
Pl
Qtz
Ky
Gt
Gl
KALB12
1100
2·1
69
Cpx
Zo
Pl
Qtz
Ky
Gt
Gl
APD680
950
2·7
148
Cpx
Zo
Ph
Qtz
Ky
Gt
FCH5
1000
2·7
118
Cpx
Zo
Ph
Qtz
Ky
Gt
APD683
1025
2·7
149
Cpx
Zo
Ph
Ksp
Qtz
Ky
APD682
1075
2·7
126
Cpx
Zo
Qtz
Ky
Gt
Gl
FCH9
1100
2·7
118
Cpx
Zo
Qtz
Ky
Gt
Gl
APD685
1125
2·7
78
Cpx
Zo
Qtz
Ky
Gt
Gl
APD684
1150
2·7
53
Cpx
Zo
Qtz
Ky
Gt
Gl
APD664
925
3·2
145
Cpx
Zo
Ph
Qtz
Ky
Gt
APD663
1000
3·2
173
Cpx
Zo
Qtz
Ky
Gt
Gl
APD686
1050
3·2
79
Cpx
Zo
Qtz
Ky
Gt
Gl
APD680
1100
3·2
29
Cpx
Zo
Qtz
Ky
Gt
Gl
Gt
Gt
Gl
Gl
Mineral symbols from Kretz (1983).
study too. Decay of K count rates has never been
observed, so that K values reported here are uncorrected.
For Na, Patiño Douce & Harris (1998) observed that
count-rate decay is a function of glass H2O content (as
inferred from difference from 100%) and calibrated a
Na correction factor that ranged from >20% to >50%
for glasses in which the uncorrected probe totals ranged
from 97 to 91%, respectively. These same correction
factors have been applied in this study (the correction
factor was not extrapolated to glasses with totals lower
than 91%, however, but instead a uniform correction of
50% was applied to Na values in all such glasses).
It is well known that increasing f O2 increases the
stability field of Fe-bearing zoisite and epidote (e.g.
Schmidt & Thompson, 1996), and it is therefore important to know the redox conditions operative during
experiments designed to study zoisite–epidote stability.
Unfortunately, f O2 cannot be buffered to specific values
by the use of solid buffers in H2O-undersaturated experiments. However, because the cell assembly is very
much larger than the sample, the f O2 conditions generated by the cell assembly will also be acting on the
sample during the experiment. Experiments with the
same cell assembly (NaCl–MgO–C) as employed in this
295
JOURNAL OF PETROLOGY
VOLUME 43
study showed that the cell assembly imposes on the
sample an f O2 that is 1–2 log units less reducing than
that generated by the quartz–fayalite–magnetite (QFM)
solid buffer (Patiño Douce & Beard, 1994, 1995).
Approach to equilibrium
Descriptions of many previous fluid-absent melting experiments clearly show that bulk equilibrium is generally
not reached [see summary by Skjerlie & Johnston (1996)
and references therein]. Non-equilibrium features in dehydration-melting experiments include neoformed
plagioclase and garnet mantling residual cores of these
phases. However, the neoformed phases are generally of
homogeneous composition, both within single crystals
and among different crystals, suggesting that they have
approached equilibrium. Disequilibrium features in the
present study include growth of new garnet and clinopyroxene on relict crystals, and persistence of corroded
cores of zoisite and kyanite surrounded by neoformed
plagioclase mantles. These mantles may have prevented
the complete breakdown of zoisite and kyanite, as will
be discussed below. The neoformed rims are generally
euhedral and homogeneous in composition, and the
glasses and neoformed phases show systematic compositional variations with pressure and temperature (see
Figs 3, 5 and 6 below). Thus, despite the presence of
disequilibrium features, we argue that the neoformed
phases represent near-equilibrium assemblages.
NUMBER 2
FEBRUARY 2002
euhedral neoformed plagioclase, clinopyroxene and kyanite form (Fig. 2f ). Relict cores of zoisite are present
even at 1100°C.
Quartz is absent at 1 GPa (Fig. 2a). It is present,
but corroded in all higher-pressure supersolidus runs.
Phengite is never present inside the plagioclase stability
field (P <2·7 GPa), within which plagioclase becomes
more abundant with decreasing pressure. Kyanite, which
is present in the starting material (Table 1), forms above
1·5 GPa (Figs 1 and 2e–h), but is corroded at 1·5 GPa
(Fig. 2b) and absent at 1·0 GPa (Fig. 2a). At 2·7 GPa
the starting material is unreacted at 1000°C. At 1025°C
phengite is absent but pseudomorphed by potassium
feldspar (Fig. 2g) whereas zoisite is slightly corroded. Our
highest temperature experiment at 2·7 GPa (1125°C)
contains corroded zoisite included in glass pools that also
contain crystals of euhedral kyanite and clinopyroxene
(Fig. 2h). At 3·2 GPa the starting material is unreacted
at 925°C. At 1000°C phengite is absent and zoisite
appears unreacted (Fig. 2i). Small pools of melt are
associated with neoformed garnet and kyanite, replacing
phengite. At 3·2 GPa and higher temperatures zoisite is
corroded or absent and pools of melt contain kyanite
and clinopyroxene. Garnet forms at 2·7 and 3·2 GPa
(Fig. 2j) and becomes more abundant with increasing
pressure. Garnet appears to behave as an inert phase at
2·1, 1·8 and 1·5 GPa, but its abundance decreases from
1·5 to 1·0 GPa.
Experimental glass compositions
EXPERIMENTAL RESULTS
Experimental conditions and phase assemblages in the
experimental products are listed in Table 2 and shown
together with the approximate location of the solidus
and other phase boundaries in Fig. 1. Zoisite appears
unreacted on the high-P, low-T side of the shaded area,
and phengite is unreacted on the high-P, low-T side of
the dotted line (Fig. 1). The solidus is located on the
basis of the absence of glass in lower-temperature experiments, but minute amounts of glass could have gone
undetected in some of these low-temperature experiments
(see below). Within the shaded area zoisite is corroded
and mantled by various phases that change with P and
T. At temperatures lower than the inferred solidus and
P <2·1 GPa, corroded zoisite crystals are surrounded by
thin mantles consisting of plagioclase, with minor
amounts of potassium feldspar and tiny needles of kyanite
(Fig. 2c). If subsolidus dehydration of zoisite occurred at
these conditions then some melting should have taken
place, but, if so, the glasses have remained undetected
in the experimental products. At 2·1 GPa and 975°C
the mantles include minor amounts of glass (Fig. 2d). As
temperature rises the abundance of glass increases and
High-pressure melts produced experimentally from the
eclogite have SiO2 contents generally >70 wt % on an
H2O-free basis (Table 3, Fig. 3). Less silicic melts (SiO2
66–67 wt %) are produced only at 1·0 GPa (Table
3, Fig. 3). The compositions of the initial melts vary
systematically with pressure (Fig. 3, Fig. 4a), from tonalites
at 1·0 GPa to tonalite–trondhjemites at 1·5 GPa, adamellites at 2·1 and 2·7 GPa, and to true granites at 3·2
GPa. This trend reflects the changing roles of zoisite and
phengite in the initial melting of the eclogite (Fig. 1). As
temperature rises above the solidus and both hydrous
phases break down, melts converge towards tonalitic–
granodioritic compositions at all pressures.
Concentrations of ferromagnesian components decrease with increasing pressure (Figs 3 and 4b) This is in
agreement with the behaviour observed in dehydrationmelting experiments of other bulk compositions, and is
a consequence of the fixed H2O budget of dehydrationmelting coupled to the increase in water solubility with
increasing pressure (e.g. Patiño Douce & Beard, 1995,
1996; Patiño Douce & McCarthy, 1998; Patiño Douce,
1999). Melts with >5 wt % FeO + MgO + TiO2 are
produced only at 1·0 GPa (Fig. 3). At greater pressures
296
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
Fig. 1. Pressure–temperature diagram showing the experimental results on the Verpenesset eclogite. Χ, experiments with glass present; Β,
charges where glass could not be confirmed. Zoisite is unreacted to the left of the shaded field, corroded in the shaded field, and absent to the
right of the shaded field. The phengite-out boundary is represented by the bold dotted line, which coincides with the solidus at pressures >2·1
GPa. Kyanite is absent at pressures lower than those limited by the line denoted kyanite out, and plagioclase is absent at pressures higher than
those limited by the line denoted plagioclase out.
FeO + MgO + TiO2 contents are generally <3 wt %,
and as low as 1–2 wt % at 2·1, 2·7 and 3·2 GPa (Table
3, Fig. 3).
Incipient melting at 2·1–3·2 GPa produces potassic
leucogranite melts (Fig. 4a, Table 3), reflecting the fact
that phengite is the dominant hydrous phase involved in
their production. However, only traces of these K-rich
melts are formed, owing to the low phengite content of
the starting material. Higher-temperature melts at P =
2·7 GPa become enriched in Ca, reflecting zoisite breakdown, but they remain leucocratic (<3 wt % FeO +
MgO + TiO2) even up to the maximum temperatures
investigated (1150°C, Fig. 3). It is interesting to note that
these remarkably leucocratic high-temperature melts are
produced from a protolith of basaltic bulk composition
(Table 1). This observation has implications for processes
at subduction zones, which are discussed below.
generally increase with increasing pressure, but they are
almost always lower than in cpx in the starting material,
except at 3·2 GPa (Fig. 5b). Na depletion relative to cpx
in the starting material is particularly strong at 1 and 1·5
GPa, reflecting abundant plagioclase crystallization. Al
contents in cpx in 1 and 1·5 GPa experiments are lower
than in cpx in the starting material, whereas Al contents
at 2·7 and 3·2 GPa are generally higher than in the starting
material (Fig. 5b). Al contents in 1·8 and 2·1 GPa runs
are comparable with those in the starting material. In most
cases, Na and Al contents in cpx decrease with rising
temperature, probably reflecting progressive incorporation
of normative jadeite into the melts. This behaviour was also
observed by Skjerlie & Johnston (1996) in their dehydrationmelting experiments on a greywacke. The behaviour of
cpx at 2·7 GPa does not follow this general trend, however,
as at this pressure Na decreases but Al increases with
rising temperature (Fig. 5b). The reason for this different
behaviour is not clear.
Clinopyroxene
Clinopyroxene is present in all run products as euhedral
and generally unzoned tabular crystals. Cpx present in 1
and 1·5 GPa runs is augite whereas at all higher pressures
it is omphacite (Table 4, Fig. 5a). Na contents in cpx
Orthopyroxene
297
Orthopyroxene is not present in the starting material but
is found as small acicular crystals in experiments at 1
JOURNAL OF PETROLOGY
VOLUME 43
298
NUMBER 2
FEBRUARY 2002
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
GPa, 950–1050°C. Opx in these experiments is an En–Fs
solid solution (Table 4, Fig. 5a), with Al2O3 contents
increasing regularly from >1 wt % at 950°C to >6 wt %
at 1050°C.
experiments at all pressures are generally indistinguishable from those of garnet in the starting material (Table 7, Fig. 7). These textural and compositional
characteristics suggest that garnet behaves as a largely
inert phase during partial melting of eclogites.
Plagioclase
Plagioclase is not present in the starting material but is
formed in experiments at 1, 1·5, 1·8 and 2·1 GPa. It
occurs as subhedral unzoned grains, commonly larger
than 10 m across. At constant temperature plagioclase
generally becomes more sodic with increasing pressure
and at all pressures it becomes more calcic with rising
temperature (Table 5, Fig. 6). These compositional trends
are consistent with the greater pressure stability of albite
relative to anorthite and with the preferential dissolution
in the melts of albite relative to anorthite as temperature
rises and melt fraction increases. Orthoclase contents in
plagioclase are always low (<10 mol %), reflecting the
low bulk K2O content.
Zoisite
Zoisite is present at 1·5–3·2 GPa but not at 1 GPa. Its
behaviour and textural relationships are complex and
are discussed in detail below, within the context of the
melting relations. It is always near end-member zoisite,
containing <2 wt % Fe2O3 (Table 6). Its composition is
essentially indistinguishable from that of zoisite present
in the starting material.
Sub-solidus and melting reactions
Figure 8 shows the P–T locations of various dehydration
reactions for zoisite and muscovite in the end-member
KCASH system [plotted with the software package
TWQ , which employs the thermodynamic database of
Berman (1988) and various updates]. Also shown is the
water-saturated basalt solidus determined experimentally
by Lambert & Wyllie (1972). The melt-absent reactions
have been plotted to attempt to better understand the
phase relationships. The dehydration reactions are metastable relative to melting reactions at temperatures higher
than the H2O-saturated solidus. However, in a fluidabsent situation, melt will be the most water-rich phase,
and will occupy a chemographic position analogous to
water [see also Vielzeuf & Montel (1994) and Skjerlie &
Johnston (1996)]. Na is not considered in the sub-solidus
reactions. The presence of Na in our natural starting
material (chiefly as jadeite in omphacitic clinopyroxene)
is likely to lower the melting temperatures and increase
the melt fraction relative to the Na-free system, but
should not affect the phase relations in any major way.
The dehydration reaction
Qz + 4Zo = Gr + 5An + 2H2O
intersects the water-saturated solidus at >650°C at
>700 MPa (Fig. 8). This intersection constrains the
minimum pressure at which dehydration-melting of zoisite is possible in quartz-bearing eclogites such as our
natural starting material, because dilution of grossular and
anorthite by other garnet and plagioclase components,
respectively, shifts the dehydration reaction to higher
pressure. The presence of muscovite in the starting material has the same effect, by rendering the following
reaction stable (see Fig. 8):
Garnet
Garnet is present in all melting experiments, but shows
clear signs of recrystallization only at 2·7 and 3·2 GPa.
Even at these pressures, recrystallization is limited to the
formation of narrow euhedral rims, generally <5 m
thick. These rims do not yield consistent chemical compositions, probably owing to inclusion in the electron
beam excitation volume of original garnet underlying
the neoformed rims. At 1 GPa garnet appears notably
corroded, emphasizing its role in orthopyroxene-forming
reactions (see below). Garnet compositions in the melting
2 zoisite + 2 quartz + muscovite = 4 anorthite +
sanidine + 2H2O.
(1)
Fig. 2. Back-scattered electron micrographs of selected experiments: (a) 950°C and 1 GPa (note absence of zoisite and quartz and presence of
orthopyroxene and abundant plagioclase); (b) 950°C and 1·5 GPa (note corroded kyanite mantled by plagioclase); (c) 900°C and 2·1 GPa (note
the very thin mantles surrounding the zoisite in the middle of the photograph; these rims consist of plagioclase, potassium-rich alkali feldspar
and kyanite in decreasing amount); (d) 975°C and 2·1 GPa (note that the largest zoisite crystal is surrounded by glass and plagioclase); (e) 1050°C
and 2·1 GPa, corroded zoisite is still present; (f ) 1100°C and 2·1 GPa (note that glass pools contain plagioclase and euhedral neoformed kyanite;
zoisite is absent); (g) 1025°C and 2·7 GPa, long pseudomorph of K-spar with kyanite crystals after phengite [also note recrystallization of
clinopyroxene (lighter colour) around the pseudomorph]; (h) 1125°C and 2·7 GPa, corroded zoisite surrounded by glass (note kyanite and
neoformed clinopyroxene); (i) 1000°C and 3·2 GPa, zoisite is unreacted; ( j) 1050°C and 3·2 GPa (note new garnet growing on the old garnet
grains).
299
1
950
T (°C):
300
4·44
0·92
17·48
2·23
0·94
0·06
4·09
3·16
0·95
Al2O3
FeO∗
MgO
MnO
CaO
Na2O
91·21
36·9
3·8
5·6
26·7
20·3
Total†
Q
C
Or
Ab
An
16·4
32·1
12·2
2·6
26·6
93·45
100·00
2·07
3·79
3·30
0·09
1·19
3·71
17·08
0·89
67·88
1·5
12·8
35·3
8·5
2·6
37·7
94·58
100·00
1·43
4·17
2·58
0·03
0·81
0·81
15·67
0·28
74·21
950
1·5
14·6
31·1
5·6
4·2
39·6
93·87
100·00
0·95
3·67
2·94
0·03
1·01
1·23
16·62
0·55
73·00
1000
1·5
12·7
33·5
12·7
2·3
31·8
94·65
100·00
2·15
3·96
2·55
0·06
1·02
2·30
15·76
0·76
71·43
1050
1·8
20·2
32·5
6·4
2·1
33·7
94·78
100·00
1·08
3·84
4·08
0·03
1·08
1·15
17·05
0·71
70·97
1050
2·1
12·9
29·6
8·3
3·8
41·8
93·68
100·00
3·20
3·42
2·03
0·02
0·39
0·76
16·63
0·31
73·25
975
2·1
15·2
44·1
10·6
0·6
25·8
95·15
100·00
1·79
5·21
3·07
0·07
0·68
0·93
16·72
0·40
71·13
1025
2·1
18·8
38·6
9·6
0·4
27·7
96·06
100·00
1·62
4·56
3·78
0·13
0·83
1·37
16·54
0·53
70·64
1050
2·1
20·3
35·9
6·6
1·7
31·1
97·57
100·00
1·12
4·24
4·09
0·04
0·93
0·99
17·31
0·79
70·49
1100
2·7
7·6
22·7
20·9
4·5
42·6
88·99
100·00
3·53
2·68
1·53
0·02
0·26
0·57
15·52
0·19
75·71
1025
2·7
5·8
23·7
21·7
5·1
42·1
90·34
100·00
3·67
2·80
1·17
0·02
0·20
0·56
15·80
0·13
75·65
1075
2·7
14·9
27·8
12·9
2·9
38·6
88·22
100·00
2·18
3·29
3·01
0·01
0·55
0·73
16·16
0·41
73·66
1100
2·7
15·1
30·0
13·0
2·1
36·5
91·94
100·00
2·20
3·55
3·04
0·03
0·55
0·92
15·90
0·57
73·25
1125
2·7
19·0
30·5
10·6
1·7
34·2
94·42
100·00
1·80
3·61
3·83
0·02
0·66
1·13
16·57
0·74
71·65
1150
3·2
8·0
26·2
30·3
1·3
32·2
91·49
100·00
5·13
3·10
1·61
0·03
0·31
0·57
14·93
0·30
74·00
1000
3·2
9·8
17·8
25·6
3·4
41·3
88·70
100·00
4·34
2·10
1·97
0·02
0·31
0·64
15·18
0·38
75·06
1050
3·2
23·3
29·9
9·8
0·2
31·0
93·27
100·00
1·66
3·53
4·70
0·02
0·96
1·65
16·36
1·24
69·88
1100
NUMBER 2
24·0
37·6
5·4
1·0
22·2
94·42
100·00
4·83
0·02
1·48
3·20
1
1000
VOLUME 43
Reported analyses are averages of 5–10 analyses of different glass pools. Q , C, Or, Ab, An—normative amount of quartz, corundum, orthoclase, albite, anorthite.
∗Fe reported as FeO.
†Original probe total.
100·00
Total
K2O
0·62
0·57
TiO2
18·11
66·37
70·53
SiO2
900
P (GPa): 1
Table 3: Glass analyses
JOURNAL OF PETROLOGY
FEBRUARY 2002
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
Fig. 3. Temperature–oxide diagram for all analysed experimental glasses. Χ, 1 GPa; Β, 1·5 GPa; Ε, 1·8 GPa; Η, 2·1 GPa; Μ, 2·7 GPa;
Ο, 3·2 GPa.
In any event, Fig. 8 shows that zoisite dehydrationmelting is possible at pressures >>1·0 GPa, which are
relevant to our experimental study as well as to melting
of eclogites in nature.
The phase relations of the natural starting material
are complex owing to the presence of two hydrous phases
with very different properties (zoisite and phengite). Zoisite appears corroded over a wide temperature interval,
shown by the shaded area in Fig. 1. The composition of
the corroded zoisite crystals is indistinguishable between
experiments at different pressures and temperatures, and
is also similar to the composition of zoisite in the starting
material (Tables 1 and 6). Breakdown of zoisite over a
finite P–T interval, as observed in our experiments, is
probably not a result of solid solution in this phase as
the composition of zoisite in the various experiments
is indistinguishable. Thus, zoisite may be metastably
preserved, aided by the formation of mantles of reaction
products. Whereas plagioclase mantles are always present
at P <2·7 GPa, at 2·7 and 3·2 GPa corroded zoisite
is always in contact with glass pools of homogeneous
composition. This may suggest that, at least at high
pressure, reaction of zoisite over a wide temperature
range may be an equilibrium process.
Initial breakdown of zoisite at 2·1 GPa is manifested
by mantling of corroded zoisite by thin rims dominated
by plagioclase and potassium feldspar and tiny needles
of kyanite (Fig. 2c). No glass was detected at this pressure
and T <975°C, but glass becomes readily observable at
higher T. At the same time the size of the plagioclase
rims increases and euhedral plagioclase forms in the glass
pools, whereas potassium feldspar disappears, quartz
becomes corroded and clinopyroxene recrystallizes to a
less sodic variety. As the pressure decreases to 1·5 GPa,
the abundance of plagioclase increases and kyanite
changes from being a product of the incongruent melting
reaction to becoming a reactant (manifested by its corroded appearance and the absence of neoformed euhedral
crystals). At 2·7 and 3·2 GPa, plagioclase is absent and
garnet is a product of the incongruent melting reaction.
These features of the experimental products constrain
the nature of the melting reactions and their changes
with pressure.
Kyanite appears corroded in all of the 1·5 GPa run
products, and is absent from all of the 1·0 GPa runs.
Neoformed plagioclase is abundant at 1·0 and 1·5 GPa.
These observations suggest that at P <1·8 GPa zoisite
breakdown follows the reaction (see Fig. 8)
301
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 2
FEBRUARY 2002
Fig. 4. (a) Normative Ab–An–Or diagram after O’Connor (1965). The dashed line traces the initial melts that form at each of the studied
pressures. The shaded field encompasses glasses experimentally produced by amphibolite dehydration-melting. Symbols as in Fig. 3. (b) All glass
analyses plotted in a ternary (Na2O + K2O)–CaO–(FeO + MgO + TiO2) diagram. The dashed line traces the initial melts that form at each
of the studied pressures. (Note that the melts become poorer in ferromagnesian components with increasing pressure.)
2 zoisite+kyanite+quartz=4 anorthite+H2O.
(2)
During zoisite dehydration-melting at these relatively low
pressures some of the anorthite dissolves in the melt and
the rest crystallizes as plagioclase.
The melt pools contain kyanite in the experimental
charges produced at P >1·5 GPa, suggesting that the
zoisite dehydration-melting reaction at these high-pressure conditions produces kyanite as a peritectic phase.
No kyanite-forming dehydration reactions are stable in
Fig. 8, suggesting that incongruent breakdown of zoisite
at high pressure could be modelled by means of the
reaction
2 zoisite=3 kyanite+3 quartz+4 CaO+H2O
(3)
where CaO and H2O represent components that are
either dissolved in the melt phase or, in the case of CaO,
could also in part go to forming other Ca-bearing phases,
such as plagioclase, garnet and diopside component of
clinopyroxene (see below). The observation that quartz
in the run products always appears corroded shows that
the SiO2 liberated by zoisite breakdown reaction (3)
also enters the melt. Both plagioclase and kyanite are
neoformed phases at 1·8 and 2·1 GPa, suggesting that
in the bulk composition that we studied zoisite breakdown
reactions (2) and (3) overlap over this pressure range.
Clinopyroxene and garnet are present in all the run
products. Garnet shows virtually no textural indications
of reaction from 1·5 to 2·1 GPa. At greater pressures
302
51·51
0·74
7·45
13·99
2·63
22·19
1·49
100·00
1·87
0·02
0·32
0·76
0·08
0·86
0·10
4·01
SiO2
TiO2
Al2O3
MgO
FeO∗
CaO
Na2O
Total
Si
Ti
Al
Mg
Fe2+
Ca
Na
Total
51·80
0·40
5·12
14·07
7·40
19·90
0·83
99·54
1·91
0·01
0·22
0·77
0·23
0·79
0·06
4·00
1·0
950
303
50·62
0·70
11·50
10·87
3·03
19·04
3·89
99·65
1·83
0·02
0·49
0·59
0·09
0·74
0·27
4·04
SiO2
TiO2
Al2O3
MgO
FeO∗
CaO
Na2O
Total
Si
Ti
Al
Mg
Fe2+
Ca
Na
Total
52·90
0·57
7·39
12·02
5·05
19·14
2·90
99·97
1·92
0·02
0·32
0·65
0·15
0·74
0·20
4·01
2·1
1000
52·45
0·91
8·86
11·20
3·60
19·88
2·96
99·86
1·90
0·02
0·38
0·60
0·11
0·77
0·21
3·99
2·1
1025
51·20
0·87
4·05
14·54
8·50
19·59
0·49
99·24
1·91
0·02
0·18
0·81
0·26
0·78
0·04
4·00
1·0
1000
50·97
0·78
12·11
11·39
3·22
19·31
2·50
100·28
1·83
0·02
0·51
0·61
0·10
0·74
0·17
3·98
2·1
1050
51·92
0·17
2·08
16·98
8·62
18·72
0·55
99·05
1·94
0·00
0·09
0·94
0·27
0·75
0·04
4·03
1·0
1050
51·69
0·92
10·06
11·08
2·96
20·48
2·53
99·71
1·87
0·03
0·43
0·60
0·09
0·79
0·18
3·98
2·1
1100
51·07
0·53
5·41
13·06
8·96
18·70
1·70
99·43
1·90
0·01
0·24
0·72
0·28
0·75
0·12
4·03
1·5
875
52·44
0·71
11·61
11·70
1·95
18·66
3·12
100·20
1·87
0·02
0·49
0·62
0·06
0·71
0·22
3·98
2·7
950
51·76
0·52
7·28
14·35
2·86
21·16
1·59
99·52
1·88
0·01
0·31
0·78
0·09
0·82
0·11
4·01
1·5
900
55·45
0·02
8·70
11·87
2·47
17·44
4·62
100·57
1·97
0·00
0·36
0·63
0·07
0·66
0·32
4·01
2·7
1000
52·37
0·43
5·30
15·23
3·38
21·29
1·43
99·43
1·91
0·01
0·23
0·83
0·10
0·83
0·10
4·01
1·5
950
Oxide values in weight percent. All values are averages of 4–5 analyses.
∗Fe reported as FeO.
2·1
975
P (GPa):
T (°C):
Clinopyroxene
1·0
900
P (GPa):
T (°C):
Clinopyroxene
52·63
0·50
13·24
9·29
2·54
16·73
3·84
98·77
1·89
0·01
0·56
0·50
0·08
0·64
0·27
3·95
2·7
1025
52·39
0·46
6·42
15·12
3·85
20·05
1·33
99·62
1·90
0·01
0·27
0·82
0·12
0·78
0·09
4·00
1·5
1000
52·43
0·68
13·53
9·51
2·34
17·23
3·41
99·13
1·88
0·02
0·57
0·51
0·07
0·66
0·24
3·94
2·7
1100
52·31
0·14
5·85
14·56
5·01
20·37
1·24
99·49
1·91
0·00
0·25
0·79
0·15
0·80
0·09
4·00
1·5
1050
50·89
0·77
14·52
9·51
2·12
17·96
2·99
98·75
1·83
0·02
0·62
0·51
0·06
0·69
0·21
3·94
2·7
1125
55·07
0·09
9·40
11·70
2·34
17·43
4·32
100·35
1·95
0·00
0·39
0·62
0·07
0·66
0·30
4·00
1·8
900
Table 4: Clinopyroxene compositions (no added H2O) and orthopyroxene compositions
48·62
0·66
19·01
6·83
2·94
18·32
2·85
99·21
1·75
0·02
0·81
0·37
0·09
0·71
0·20
3·93
2·7
1150
56·08
0·01
8·68
11·69
1·84
17·35
4·74
100·39
1·98
0·00
0·36
0·62
0·05
0·66
0·32
4·00
1·8
950
53·55
0·43
12·06
10·12
2·15
15·99
4·34
98·63
1·92
0·01
0·51
0·54
0·06
0·61
0·30
3·96
3·2
1000
55·42
0·06
8·84
11·94
2·17
17·91
4·30
100·63
1·96
0·00
0·37
0·63
0·06
0·68
0·30
4·00
1·8
1050
54·78
0·03
8·57
11·83
2·12
17·45
4·84
99·62
1·96
0·00
0·36
0·63
0·06
0·67
0·34
4·02
3·2
1100
55·05
0·08
8·45
11·64
2·09
17·71
4·25
99·27
1·97
0·00
0·36
0·62
0·06
0·68
0·30
3·99
2·1
800
55·04
0·07
8·23
11·85
1·96
17·89
4·41
99·45
1·97
0·00
0·35
0·63
0·06
0·69
0·31
4·01
2·1
900
53·72
0·17
1·04
24·77
17·49
1·85
0·10
99·14
1·98
0·00
0·05
1·36
0·54
0·07
0·01
4·00
1·0
950
53·13
0·21
3·68
25·35
15·53
1·67
0·11
99·68
1·92
0·01
0·16
1·37
0·47
0·06
0·01
4·00
1·0
1000
Orthopyroxene
54·86
0·08
9·00
11·51
2·13
17·00
4·48
99·06
1·97
0·00
0·38
0·62
0·06
0·65
0·31
4·00
2·1
850
50·05
0·30
6·09
23·00
18·71
1·58
0·11
99·84
1·84
0·01
0·26
1·26
0·58
0·06
0·01
4·02
1·0
1050
56·40
0·05
5·83
11·71
2·90
17·76
4·85
99·49
2·03
0·00
0·25
0·63
0·09
0·68
0·34
4·02
2·1
950
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
JOURNAL OF PETROLOGY
VOLUME 43
NUMBER 2
FEBRUARY 2002
Fig. 5. (a) Pyroxene compositions in eclogite melting experiments [classification scheme after Morimoto (1988)]. Composition of clinopyroxene
in the starting material is shown with a large filled diamond. The arrows show general compositional trends with rising temperature at 1·0–1·5,
2·1 and 2·7 GPa. Orthopyroxene (present only at 1·0 GPa) plots in the enstatite–ferrosilite (En–Fs) field. (b) Al and Na contents in clinopyroxene
as a function of temperature and pressure. Dashed lines show composition of clinopyroxene in the starting material.
garnet is clearly formed in the experiments, and its
abundance increases from 2·7 to 3·2 GPa. Garnet
formation probably results from the following dehydration-melting reaction suggested by Boettcher (1970):
ferromagnesian content of phengite forms garnet by an
incongruent dehydration-melting reaction such as
zoisite + quartz = grossular + kyanite + melt. (4)
Except at 3·2 GPa, clinopyroxene in all glass-bearing
run products is less sodic than that in the starting material,
showing that the albite component of the melts is largely
supplied by breakdown of the jadeite component of
clinopyroxene. Formation of a less sodic clinopyroxene
can be represented by the coupled substitution NaAlcpx
↔ CaMgcpx, which suggests that Ca liberated by zoisite
The experimental charge at 1000°C and 3·2 GPa contains
unreacted zoisite whereas phengite has broken down to
garnet, kyanite and glass. This, together with the notable
backbending of the phengite reaction boundary between
2·7 and 3·2 GPa, suggests that at high pressures the
phengite + quartz = kyanite + garnet + melt. (5)
304
950
9·35
Na2O
305
5·00
Total
0·56
4·97
0·02
1·0
5·00
0·00
0·42
0·57
0·02
1·58
2·41
100·56
0·05
4·80
11·80
0·45
29·87
53·59
1000
1·0
5·01
0·00
0·02
0·99
0·02
1·97
2·01
99·40
0·00
0·19
19·92
0·41
35·90
42·98
1050
1·5
5·00
0·02
0·59
0·39
0·08
1·39
2·59
101·37
0·28
6·89
8·28
0·46
26·78
58·69
875
1·5
4·980
0·02
0·57
0·37
0·01
1·39
2·62
99·66
0·26
6·57
7·79
0·23
26·30
58·51
900
1·5
4·98
0·01
0·61
0·35
0·01
1·36
2·65
99·72
0·17
7·02
7·21
0·34
25·78
59·20
950
1·5
4·97
0·01
0·57
0·38
0·01
1·36
2·64
101·15
0·15
6·72
7·98
0·28
26·25
59·77
1000
Oxide values in weight percent. All values are averages of 4–5 analyses.
∗Fe reported as FeO.
0·81
0·01
Na
K
0·37
0·02
0·04
0·23
Fe
Ca
1·37
1·04
Al
2·63
2·87
100·14
0·34
6·51
7·75
0·45
26·15
Si
99·66
4·91
CaO
Total
0·97
FeO∗
0·20
19·86
K 2O
64·37
SiO2
Al2O3
58·95
1·0
T (°C):
900
P (GPa): 1·0
Table 5: Plagioclase compositions, no added H2O
1·5
5·05
0·00
0·06
0·97
0·01
2·06
1·95
99·59
0·03
0·66
19·47
0·34
37·41
41·68
1050
1·8
4·96
0·02
0·60
0·32
0·01
1·33
2·68
100·83
0·30
6·99
6·84
0·24
25·57
60·89
900
1·8
4·97
0·02
0·58
0·37
0·01
1·34
2·65
100·30
0·43
6·72
7·75
0·20
25·54
59·66
950
1·8
4·98
0·01
0·54
0·43
0·02
1·39
2·59
99·65
0·15
6·19
8·93
0·49
26·22
57·67
1050
2·1
4·97
0·10
0·63
0·24
0·02
1·20
2·78
99·99
1·73
7·31
4·96
0·66
22·87
62·46
900
2·1
4·97
0·11
0·59
0·25
0·02
1·27
2·74
100·20
1·89
6·85
5·24
0·49
24·15
61·58
950
2·1
4·97
0·03
0·63
0·30
0·02
1·29
2·71
100·49
0·57
7·31
6·39
0·46
24·68
61·08
975
2·1
4·98
0·08
0·58
0·30
0·01
1·30
2·70
100·32
1·47
6·71
6·30
0·35
24·80
60·69
1000
2·1
5·042
0·014
0·642
0·384
0·007
1·426
2·570
99·87
0·24
7·36
7·97
0·18
26·92
57·19
1025
2·1
4·989
0·014
0·591
0·382
0·010
1·368
2·624
101·12
0·24
6·90
8·07
0·26
26·27
59·38
1050
2·1
5·034
0·007
0·510
0·511
0·010
1·554
2·442
99·73
0·13
5·80
10·53
0·27
29·10
53·90
1100
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
8·031
8·054
8·046
8·053
8·059
7·988
8·024
8·026
Oxide values in weight percent. All values are averages of 4–5 analyses.
∗Fe reported as FeO.
8·018
8·015
8·024
8·071
8·017
8·016
7·993
8·041
Total
8·023
2·055
0·070
0·072
2·097
2·061
0·098
0·100
2·108
2·100
0·094
0·062
1·959
2·048
0·073
0·112
2·037
2·043
0·095
0·070
2·027
2·054
0·060
0·057
2·113
2·017
0·122
0·063
2·005
1·977
0·089
0·063
Ca
2·036
0·112
2·059
Fe3+
2·956
2·951
2·955
2·966
2·898
2·945
2·975
2·931
2·918
2·897
2·942
2·936
3·002
2·919
2·996
2·937
2·958
2·939
Al
97·45
97·80
2·929
2·921
97·51
98·67
2·946
2·920
97·82
98·00
2·992
2·973
97·18
97·09
2·959
2·982
97·38
96·42
2·977
2·974
96·97
96·90
2·899
2·958
98·45
97·23
2·952
2·990
97·53
97·80
2·966
2·930
97·54
Total
24·94
25·06
Si
1·10
25·05
25·59
25·08
25·91
25·59
24·15
24·90
24·71
24·89
24·49
24·94
25·52
24·83
24·45
32·70
CaO
24·21
1·14
1·54
1·59
1·49
0·99
1·15
1·75
1·50
1·09
0·95
0·89
1·95
0·99
1·41
1·77
FeO∗
0·99
38·60
38·29
32·78
32·81
38·08
38·79
32·38
32·62
38·12
39·52
33·34
32·40
38·73
38·45
32·18
32·08
38·91
38·53
32·31
32·40
38·68
37·52
32·97
32·66
39·01
38·57
33·22
32·69
32·51
39·22
38·93
38·20
SiO2
Al2O3
32·94
2·7
2·1
1100
1050
2·1
2·1
1025
1000
2·1
2·1
975
950
900
850
2·1
2·1
1·8
1·8
1·8
1·5
1·5
900
1·5
800
2·1
306
1050
2·1
(8)
1·5
zoisite + NaAlclinopyrexene + quartz = melt +
kyanite + CaMgclinopyroxene ± garnet.
875
At high pressures zoisite dehydration-melting probably
takes place by a reaction such as
P (GPa):
zoisite + NaAlclinopyroxene + kyanite + quartz ± garnet
= melt + plagioclase + CaMgclinopyroxene
± orthopyroxene.
(7)
FEBRUARY 2002
T (°C):
(see also Patiño Douce & McCarthy, 1998). It should be
noted that quartz is always absent at 1·0 GPa, but is
present in all higher-pressure experimental products.
Reaction (6) is also compatible with the fact that there is
less garnet at 1 GPa than in higher-pressure experiments.
Summarizing and integrating all of these observations,
we conclude that dehydration-melting of zoisite-bearing
eclogites at pressures lower than >1·8 GPa takes place
by the following melting reaction:
Table 6: Zoisite compositions, no added H2O
clinopyroxene + garnet + quartz = orthopyroxene +
plagioclase
(6)
950
breakdown also enters clinopyroxene in addition to entering the melt and forming anorthite or grossularite.
Orthopyroxene is present only at 1·0 GPa. Because
the equilibration pressure of the natural starting material
is greater than this (>2·5 GPa), orthopyroxene in the
run products probably formed by the decompression
reaction
900
Fig. 6. Plagioclase compositions in eclogite melting experiments as a
function of temperature and pressure. The starting material contains
no plagioclase.
1000
NUMBER 2
1000
VOLUME 43
950
JOURNAL OF PETROLOGY
307
38·87
0·12
22·72
0·02
10·34
17·32
9·17
0·80
99·35
2·94
0·01
2·03
0·00
1·17
1·10
0·74
0·05
8·04
SiO2
TiO2
Al2O3
Cr2O3
MgO
FeO∗
CaO
MnO
Total
Si
Ti
Al
Cr
Mg
Fe
Ca
Mn
Total
39·08
0·06
23·54
0·04
11·91
17·00
8·54
0·51
100·67
2·90
0·00
2·06
0·00
1·32
1·06
0·68
0·03
8·06
2·1
1000
38·95
0·02
23·45
0·00
11·65
17·51
7·85
0·47
99·90
2·92
0·00
2·07
0·00
1·30
1·10
0·63
0·03
8·05
1·0
950
39·26
0·09
22·01
0·03
10·25
17·93
10·89
0·29
100·74
2·95
0·00
1·95
0·00
1·15
1·13
0·88
0·02
8·07
2·1
1025
40·38
0·01
23·27
0·00
13·98
16·01
7·63
0·25
101·53
2·95
0·00
2·00
0·00
1·52
0·98
0·60
0·02
8·05
1·0
1000
38·87
0·01
22·54
0·11
11·74
16·92
9·37
0·33
99·89
2·92
0·00
2·00
0·01
1·32
1·06
0·75
0·02
8·08
2·1
1050
38·46
0·07
22·11
0·04
7·48
20·67
9·95
1·11
99·89
2·95
0·00
2·00
0·00
0·86
1·33
0·82
0·07
8·04
1·0
1050
39·28
0·10
22·53
0·05
11·89
16·57
8·24
0·38
99·04
2·96
0·01
2·00
0·00
1·34
1·04
0·66
0·02
8·03
2·1
1100
39·10
0·03
22·83
0·09
11·17
16·10
10·10
0·50
99·92
2·93
0·00
2·02
0·01
1·25
1·01
0·81
0·03
8·06
1·5
875
39·54
0·00
22·95
0·00
11·87
17·72
7·85
0·26
100·19
2·95
0·00
2·02
0·00
1·32
1·11
0·63
0·02
8·04
2·7
950
39·32
0·05
22·97
0·00
11·17
16·39
9·11
0·22
99·24
2·96
0·00
2·03
0·00
1·25
1·03
0·73
0·01
8·02
1·5
900
38·46
0·06
23·76
0·02
9·37
18·21
10·36
0·70
100·94
2·89
0·00
2·10
0·00
1·05
1·14
0·83
0·04
8·06
2·7
1000
40·15
0·06
23·21
0·11
12·77
16·64
6·53
0·41
99·88
2·98
0·00
2·03
0·01
1·41
1·03
0·52
0·03
8·00
1·5
950
Oxide values in weight percent. All values are averages of 4–5 analyses.
∗Fe reported as FeO.
2·1
975
38·45
0·05
22·07
0·13
9·71
20·02
9·43
0·32
100·18
2·93
0·00
1·98
0·01
1·10
1·27
0·77
0·02
8·08
SiO2
TiO2
Al2O3
Cr2O3
MgO
FeO∗
CaO
MnO
Total
Si
Ti
Al
Cr
Mg
Fe
Ca
Mn
Total
P (GPa):
T (°C):
1·0
900
P (GPa):
T (°C):
Table 7: Garnet compositions, no added H2O
39·24
0·00
22·09
0·14
9·56
17·82
10·56
0·66
100·08
2·97
0·00
1·97
0·01
1·08
1·13
0·86
0·04
8·05
3·2
1100
38·97
0·00
23·06
0·09
12·46
17·57
8·10
0·17
100·42
2·91
0·00
2·03
0·01
1·39
1·10
0·65
0·01
8·08
1·5
1000
39·23
0·06
23·23
0·17
12·97
16·60
7·40
0·20
99·86
2·92
0·00
2·04
0·01
1·44
1·03
0·59
0·01
8·05
1·5
1050
39·67
0·06
22·67
0·02
11·56
16·79
7·98
0·21
98·96
2·98
0·00
2·01
0·00
1·30
1·06
0·64
0·01
8·01
1·8
900
38·68
0·11
22·23
0·02
9·08
18·62
9·98
0·36
99·07
2·96
0·01
2·00
0·00
1·04
1·19
0·82
0·02
8·03
1·8
950
39·63
0·05
22·24
0·16
10·26
17·63
9·34
0·31
99·61
2·99
0·00
1·98
0·01
1·15
1·11
0·75
0·02
8·02
1·8
1050
38·90
0·04
22·38
0·00
9·72
19·78
8·23
0·21
99·26
2·96
0·00
2·01
0·00
1·10
1·26
0·67
0·01
8·03
2·1
800
38·97
0·02
22·63
0·00
11·83
17·16
8·52
0·35
99·47
2·93
0·00
2·01
0·00
1·33
1·08
0·69
0·02
8·06
2·1
850
38·85
0·08
22·10
0·22
9·16
19·82
8·80
0·42
99·44
2·97
0·00
1·99
0·01
1·04
1·27
0·72
0·03
8·03
2·1
900
39·30
0·03
23·00
0·00
13·61
16·14
7·04
0·18
99·30
2·93
0·00
2·02
0·00
1·51
1·01
0·56
0·01
8·05
2·1
950
SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
JOURNAL OF PETROLOGY
VOLUME 43
Fig. 7. Garnet compositions in eclogite melting experiments. Compositions of garnet cores and rims in the starting material are shown
with large filled diamonds.
Transition between these two melting reactions takes
place over a fairly wide pressure interval, which in the
particular starting material that we studied is bracketed
between 1·5 and 2·1 GPa.
Breakdown of phengite supplies potassium to the melt,
and at high pressure it also leads to crystallization of
peritectic garnet [reaction (5)]. K-feldspar, which forms
as a results of muscovite dehydration-melting (e.g. Patiño
Douce & Harris, 1998) is present in near-solidus runs at
2·1 and 2·7 GPa, but not at other pressures (Table 2).
The absence of K-feldspar from many run products in
which phengite is reacting out is a reflection of the
proximity between the muscovite dehydration-melting
solidus and the kfs-out phase boundary [see Patiño Douce
& Beard (1995) and Patiño Douce & Harris (1998) for
details].
The zoisite phase boundary has a positive dP/dT slope
over the entire pressure range that we investigated. This
slope steepens above 2·1 GPa, probably as a result of
crystallization of peritectic garnet. The phengite phase
boundary undergoes a sharp backbending at P >2·7 GPa
(Fig. 1), also in response to stabilization of peritectic
garnet. As a result of this backbending, phengite reacts
out completely at P = 3·2 GPa at a temperature lower
than that of the onset of zoisite dehydration-melting.
Phengite is thus the only hydrous phase that contributes
to the initial melt formed at 3·2 GPa. This behaviour is
distinctly different from that observed at lower pressures,
at which both zoisite and phengite contribute, albeit in
different proportions, to the initial melts.
The phase relationships at pressures >1 GPa, described
above and interpreted from the experimental run products, can be graphically viewed in the stereoscopic
NUMBER 2
FEBRUARY 2002
tetrahedral projection of Fig. 9a–c. An Excel routine for
plotting stereoscopic projections was made, using the
algorithm of Spear (1980). All experimental charges at
pressures >1 GPa contain quartz so that the phases can
be projected through this phase. The phases are projected
from quartz into the tetrahedron CN (CaO + Na2O)–
H2O–AK (Al2O3 + K2O)–FM (FeO + MgO + MnO).
TiO2 has been omitted (which in principle is equivalent
to projecting through rutile), but this does not affect the
phase relationships. The approximation was made that
the water content of the phases equals 100 minus probe
total. This may not be completely true, but if the estimated
water contents are somewhat too high or low, this will
move the projection points slightly towards or away from
the H2O apex of the tetrahedron and will not alter the
conclusions. Figure 9a shows that at P = 2·1 GPa the
zoisite dehydration-melting reaction (7) is supported by
the projection, because the triangle zoisite + jadeite +
kyanite intersects the anorthite + diopside + glass
triangle. Zoisite plots slightly above the CN–AK–H2O
plane as a result of the presence of Fe2O3 in this phase.
It should be noted that the tie-line zoisite + kyanite
pierces the anorthite + diopside + glass triangle close
to the anorthite corner because zoisite and anorthite
project nearby. Because of these relationships, zoisite
dehydration-melting is bound to produce small amounts
of H2O-rich melt within the anorthite stability field.
As the pressure increases, kyanite changes from being
consumed to being produced, and the projection shows
that the only way this can be explained is by formation
of a phase more calcic than zoisite. At 2·7 and 3·2 GPa,
the garnet-forming reaction (8) is thus supported (Fig.
9b). At lower pressures growth of tiny kyanite needles is
supported by a reaction such as (3) in which CaO dissolves
in the melt. Phengite breaks down in reaction with quartz
and the jadeite component of omphacite to kyanite,
glass and a MgO-silicate that is probably MgO-rich
clinopyroxene at low pressures and pyrope component
of garnet at high pressures. These phase relationships
are displayed in Fig. 9c.
COMPARISON WITH PREVIOUS
EXPERIMENTAL STUDIES
Our experiments suggest that, under fluid-absent conditions, zoisite is stable to very high temperatures at high
pressures and that it can thus be an important water
reservoir in the deep continental crust and in subducting
oceanic crust. To our knowledge, the only previous
studies of zoisite stability under fluid-absent conditions
are those by Thompson & Ellis (1994) and Boettcher
(1970). The results of Thompson & Ellis (1994) on the
CMASH system are not consistent with our results, as
they suggest that zoisite dehydration-melting occurs only
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SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
Fig. 8. Pressure–temperature diagram showing the CKASH (CaO, K2O, Al2O3, SiO2, H2O) dehydration reactions calculated by the TWQ
software [Berman (1988) and various updates]. The reactions are plotted as continuous lines at temperatures below those of the wet basalt
solidus (Lambert & Wyllie, 1972) and as dotted lines at higher temperatures. The three subduction geotherms are from Thompson & Ellis
(1994). The phengite-out boundary and the line marking the beginning of zoisite breakdown (bold continuous lines) determined in this study
are plotted, and so is the approximate P–T field covered by eclogites in the western gneiss region (shaded oval area).
slightly above the water-saturated solidus. However, those
workers considered zoisite dehydration-melting within
the amphibole stability field; they did not discuss zoisite
dehydration-melting above the upper pressure limit of
amphibole. On the other hand, the results of Boettcher
(1970) on the CASH system are compatible with our
results. Boettcher (1970) demonstrated that the thermal
stability of zoisite expands considerably in the absence
of a hydrous fluid phase, that the zoisite dehydrationmelting reaction slopes positively in P–T space and that
it typically occurs above 1050°C. These temperatures
are >100°C higher than those that we determined
(Fig. 10), but the two dehydration-melting reactions are
approximately parallel. The lower breakdown temperature of zoisite in our study is probably the result of
the presence of Na2O and K2O in our natural starting
material.
GEOLOGICAL IMPLICATIONS
The results of this study have implications for processes
in overthickened continental crust and in subduction
zones.
Overthickened continental crust
The experiments show that high-pressure metamorphosed mafic rocks containing hydrous phases such
as zoisite and phengite are capable of producing small
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Fig. 9. Stereoscopic tetrahedral projections at P >1·0 GPa employing the algorithm of Spear (1980). All phases are projected through quartz
and rutile into the CN (CaO + Na2O)–FM (FeO + MgO + MnO)–AK (Al2O3 + K2O)–H2O tetrahedron. Quartz is absent in the 1 GPa
experiments and the 1 GPa phase relations can thus not be portrayed in the projections. The small black dots represent all the experimentally
produced glasses at P >1·0 GPa. The open circle marked 1075/2·1 GPa in (a) represents the projection position of the glass in the experiment
at the same conditions. Ε, composition of the starting material, which projects close to the tie-line Cpx–Gt: Cpx and Gt represent the composition
of clinopyroxene and garnet, respectively, in the starting material. Ph and Zo mark the projection position of phengite and zoisite, respectively,
in the starting material. Gr, grossular; Ky, kyanite; An, anorthite; Jd, jadeite; Di, diopside; Py, pyrope; Alm, almandine (all end-member
compositions).
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SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
Fig. 10. P–T diagram comparing the location of the phengite-out boundary and the beginning of zoisite breakdown determined in this study,
and the location of the zoisite + quartz dehydration-melting reaction determined experimentally by Boettcher (1970). Also compared is the
zoisite + amphibole + quartz dehydration-melting reaction proposed by Thompson & Ellis (1994) in the CMASH system. The dash–dot line
is a hypothetical decompression path characterized by significant heating causing dehydration-melting of zoisite eclogites (Hacker & Peacock,
1994; Patiño Douce & McCarthy, 1998).
amounts of felsic melts if high enough temperatures are
achieved. This has potentially important implications for
the tectonic development of orogenic belts that contain
thick sections of mafic rocks. Eclogites in such settings
may form in response to fluid-fluxing as exemplified by
the eclogites of Holsnøy north of the town of Bergen,
in the Norwegian Caledonides. Eclogitization there is
restricted to shear zones that acted as channels for H2Oand CO2-bearing fluids in otherwise dry granulitic rocks
(e.g. Jamtveit et al., 1990). This H2O becomes locked up
in the deep crustal high-grade rocks as a result of the
formation of zoisite from plagioclase. Temperatures near
the base of thickened continental crust are generally
<950°C at P >2 GPa (e.g. Patiño Douce et al., 1990),
and our results show that these temperatures are too low
to trigger zoisite dehydration-melting (compare with Fig.
1). However, build-up of major mountain chains is followed by orogenic (extensional) collapse and exhumation
of deep-seated eclogites. In this paper, orogenic collapse
describes the destabilization of overthickened crust and
formation of extensional faults, which causes unroofing
of deep-seated rocks. The explanation for orogenic
collapse is debated but one suggestion is that it is caused by
mechanical instability of the lower part of the lithospheric
mantle, e.g. delamination of part of the thickened lithospheric mantle, or convective thinning (e.g. Bird, 1979;
Houseman et al., 1981; Kay & Kay, 1993; Sun & Murrell,
1994). During delamination, cold lithospheric mantle is
replaced by hot asthenospheric mantle. Basaltic magmas
that form by decompression melting may underplate and
heat the lower crust to temperatures high enough for
dehydration-melting of high-pressure zoisite-bearing
rocks. Our experiments show that this would generate
small amounts of felsic melts, which would lower the
strength of the crust as explained above. The resulting
felsic rocks would be Sr rich and enriched in the LREE.
Zoisite-involved melting may also occur during exhumation of deep-seated ultra-high-pressure rocks. It is
clear that many deep-seated rocks are exhumed along
cooling paths, but if the rocks heat on their return to the
surface (Fig. 10; Hacker & Peacock, 1994), widespread
melting would be expected. Such a P–T–t path would
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be expected, for example, if exhumation is initiated by
delamination of the lithospheric mantle.
Subduction-zone magmatism
Zoisite forms in subducting oceanic crust from continuing
dehydration of low-T hydrous phases, so that zoisite is
probably the main hydrous phase at pressures between
2·3 and 3·5 GPa under water-saturated conditions (Poli
& Schmidt, 1997). Phengite may also be present if clayrich sediments are subducted. As shown by Boettcher
(1970) and supported by our experimental results, the
stability of zoisite seems to expand considerably under
fluid-absent conditions.
Whether or not the subducting slab melts or dehydrates
is a matter of debate, but recent studies such as that by
Stoltz et al. (1996) strongly suggest that silicic melts
form in subducting slabs. The relationships between
the conditions for zoisite breakdown and subduction
geotherms may be critical in this respect. In Fig. 8 we
show three subduction geotherms from Thompson &
Ellis (1994). Subduction of old and cold crust follows a
low-temperature–high-pressure geotherm that does not
reach the breakdown temperature of zoisite at pressures
below 5·0 GPa. Subduction along such gradients is well
within the lawsonite stability field (Poli & Schmidt, 1997).
Subduction of young and hot crust causes zoisite to
undergo either subsolidus dehydration or dehydrationmelting, and amphibole dehydration-melting would also
be possible in this case. Subduction geotherms between
these two extremes traverse P–T regimes in which zoisite
is likely be the most important hydrous phase in mafic
rocks. Rocks will not cross the zoisite subsolidus dehydration reactions along such geotherms, so that zoisite
can be expected to undergo dehydration-melting in subduction zones in which the oceanic crust is neither too
cold nor too hot. Phengite dehydration-melting could
also take place if clay-rich sediments are subducted along
moderate geotherms.
Our experiments show that small melt fractions are
produced by fluid-absent zoisite dehydration-melting. It
is well known that zoisite can be the main carrier of Sr
in plagioclase-free rocks. As an example, analyses of
zoisite from the subduction-related Catalina schist show
large concentrations of Sr and also of LREE, Y and Pb
(e.g. Hickmott et al., 1992; Nagasaki & Enami, 1998).
Consequently, small-scale zoisite dehydration-melting in
subducted oceanic crust may be an effective way of
enriching the mantle wedge in these incompatible trace
elements. It would also reduce the solidus temperature
of the mantle wedge as a result of the addition of hydrous
melts. Subduction-related calc-alkaline basaltic rocks are
frequently enriched in Sr, and zoisite dehydration-melting
may be an important way of enriching the mantle wedge,
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FEBRUARY 2002
from which the basalts are ultimately derived, in Sr (see
also Nagasaki & Enami, 1998). Assuming a bulk-rock
abundance of 200 ppm Sr, formation of 5% melt by
zoisite dehydration-melting could produce felsic melts
with as much as 4000 ppm Sr, if one assumes a perfectly
incompatible behaviour. Phengite dehydration-melting
produces melts that are strongly potassic and rich in Rb
and Ba, and it could be instrumental in enriching the
mantle wedge in these incompatible elements.
Zoisite dehydration-melting vs amphibole
dehydration-melting in subducting slabs
The range in melt compositions obtained in previous
experimental studies of amphibolite dehydration-melting
is shown in Fig. 4a. A comparison with the melts produced
in our experiments shows that the melts produced by
zoisite dehydration-melting are indistinguishable in their
major element compositions from those produced by
amphibole dehydration-melting. However, because amphibolite dehydration-melting takes place within the
plagioclase stability field, the trace element signature of
these melts will be different from those generated by
zoisite dehydration-melting. The Sr contents of amphibolite melts will be lower, and amphibolite melts will
also have negative Eu anomalies, which the zoisite melts
will lack.
At P >2·7 GPa, only phengite is involved in initial
melting of the eclogite, and the leucogranitic melts produced in such circumstances are clearly different from
those produced by amphibole dehydration-melting (Fig.
4a).
The P–T path followed by a metamorphosed basalt
determines whether the rock undergoes dehydrationmelting from the amphibolite facies (young and hot
subduction, Fig. 8), or whether it first undergoes subsolidus dehydration into the eclogite facies and then, if
hydrous phases such as zoisite and phengite persist,
dehydration-melting from the eclogite facies when these
hydrous phases break down. Our results show that, in
most circumstances, the metamorphic P–T path will not
affect the major element compositions of potential slab
melts. However, the trace element signature will be
different.
There is one possible exception to the conclusion in
the last paragraph, and that is if eclogite melting takes
place at high pressure and under conditions such that
extraction of very small melt fractions proceeds at approximately the same rate as that at which they form
(e.g. in an active deformation regime; Sawyer, 1996). In
this case, incipient eclogite melting could give rise to
segregations of potassic leucogranite melts that would
also be enriched in incompatible elements, such as Rb
and Ba, that reside chiefly in micas. These melts could
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SKJERLIE AND PATIÑO DOUCE
ZOISITE DEHYDRATION-MELTING
form small intrusions or, perhaps more likely, could react
with mantle rocks and produce metasomatic effects. The
residual eclogite would then be strongly depleted in large
ion lithophile elements (LILE), but would still be able to
generate hydrous Sr-rich tonalitic–trondhjemitic melts
by zoisite dehydration-melting. This sequence of two
distinct melting events that yield very different products
is unlikely to take place during dehydration-melting of
amphibolites, because most of the K, LILE and H2O of
such rocks are concentrated in a single phase (amphibole).
CONCLUSIONS
Many natural eclogites contain small amounts of phengite
and zoisite. We have shown that such rocks undergo
dehydration-melting at relatively high temperatures
(>900°C), producing small fractions of silica-rich melts.
At pressures lower than >3 GPa the two hydrous phases
break down at similar conditions, so that the melts are
of tonalitic to granodioritic composition and probably
highly enriched in Sr, Ba, Rb and LREE. The stabilities
of the two hydrous phases diverge at higher pressures,
giving rise to discrete pulses of melting: Ba- and Rbrich leucogranitic melts produced when phengite breaks
down, followed by Sr-rich tonalitic–trondhjemitic melts
formed in response to zoisite breakdown. Melts formed
by dehydration-melting of eclogites are indistinguishable
in their major element compositions from those formed at
lower pressure from dehydration-melting of amphibolites,
but their trace element signatures may be distinctive,
especially with respect to those elements that are strongly
included in plagioclase (e.g. Sr, Eu). Eclogite dehydrationmelting in subducting slabs may thus be an important
factor in causing mantle wedge metasomatism.
Zoisite-bearing eclogites may be abundant in deep
levels of orogenic belts in which the thickened crust
contains mafic lithologies. Zoisite dehydration-melting in
such environments, perhaps triggered by an increase in
mantle heat flow arising from lithospheric delamination,
may perhaps aid in weakening the crust sufficiently to
lead to catastrophic orogenic collapse and exhumation
of deep-seated rocks.
ACKNOWLEDGEMENTS
Funding for this project was provided by the Norwegian
Research Council (NFR-122198/410). Critical review by
Kurt Bucher and an anonymous reviewer substantially
improved the manuscript.
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