JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 11
PAGES 2043^2062
2009
doi:10.1093/petrology/egp068
The High P^T Stability of Hydroxyl-apatite in
Natural and Simplified MORBçan
Experimental Study to 15 GPa with Implications
for Transport and Storage of Phosphorus and
Halogens in Subduction Zones
JU«RGEN KONZETT1* AND DANIEL J. FROST2
INSTITUT FU«R MINERALOGIE UND PETROGRAPHIE, UNIVERSITA«T INNSBRUCK, INNRAIN 52, A-6020 INNSBRUCK,
1
AUSTRIA
2
BAYERISCHES GEOINSTITUT, UNIVERSITA«T BAYREUTH, D-95440 BAYREUTH, GERMANY
RECEIVED OCTOBER 16, 2008; ACCEPTED SEPTEMBER 15, 2009
ADVANCE ACCESS PUBLICATION OCTOBER 12, 2009
Experiments have been conducted in the range 3^15 GPa and 850^
18008C to investigate the P^T stability field of OH-apatite in an
average mid-ocean ridge basalt (MORB) and a model Mg-basalt,
to study the compositional evolution of apatite and its breakdown
products and the partitioning of P between phosphates and silicates.
In the bulk compositions investigated OH-apatite is stable to
57·5 GPa at 9508C in a typical eclogite assemblage
garnet þ omphacite þ SiO2 þ TiO2. This is 5 GPa below the
breakdown P of pure OH-apatite. The high-P breakdown product is
tuite [-Ca3(PO4)2]. Both apatite and tuite are stable in a wide
range of subduction zone Tregimes but not along an average mantle
adiabat. This precludes apatite or tuite stability in the asthenospheric mantle. Apatite may be stable in cold continental lithosphere
(40 mW/m2) but is restricted to P54^5 GPa.The apatite breakdown reaction is an important limit for the crust^mantle transport
of Cl in subduction zones and can contribute to the Cl depletion of
subducted cust. Both apatite and tuite are important storage sites
for large ion lithophile elements (LILE) and rare earth elements
(REE), therefore apatite breakdown does not greatly affect LILE
or REE transport in subduction zones. In an eclogite assemblage
only garnet can accommodate significant P. In the presence of apatite
or tuite, P2O5 contents in garnet range from 0·2 to 0·6 wt %
between 3 and 11 GPa and increase to 0·8 wt % at 15 GPa in the
absence of a detectable phosphate phase. The P-storage capacity of
*Corresponding author. Telephone: þ43-(0)512-507-5506.
Fax: þ43-(0)512-507-2926. E-mail: [email protected]
clinopyroxene is limited to 250 ppm. Because of the extreme preference of P for the garnet structure, virtually the entire P budget of subducted MORB will be locked up in garnet well into the lower
mantle provided fO2 is high enough to prevent the stability of a
metal phase.
KEY WORDS:
apatite; MORB; tuite; phosphorus; high P^T stability
I N T RO D U C T I O N
Apatite [Ca5(PO4)3(OH, F, Cl)] is the most abundant naturally occuring phosphate and one of the most important
halogen-bearing minerals on Earth. Apatite and merrillite
[(Mg,Fe)2Ca18^x(Y,REE)xNa2^xP14O56] are also the two
most common phosphates in extraterrestrial rocks (Jolliff
et al., 2006). Apatite is not only of importance from a geological point of view but also plays a key role as a major P
source for terrestrial ecosystems (Philippelli, 2002, and
references therein). In addition, apatite is an essential constituent of many types of biomineralization involving
bone and dental tissue (Elliott, 1994, 2002). Apatite sensu
stricto is the most common representative of the apatite
group of minerals with the general formula M10(ZO4)6X2.
The Author 2009. Published by Oxford University Press. All
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JOURNAL OF PETROLOGY
VOLUME 50
Because of its remarkable flexibility the apatite structure
can accommodate a multitude of metal cations on the Mposition (e.g. Ca, Sr, Ba, Y, REE, Pb, U, Th, Zn, Cu) and
anion complexes such as (AsO4)3^, (SO4)2^, (CO3)2^,
(SiO4)4^, (VO4)3^ that replace (PO4)3^ and OH^, Cl or F
(Pan & Fleet, 2002, and references therein). In addition to
the chemical complexity, order^disorder processes on the
halogen site give rise to considerable structural complexity
involving hexagonal and monoclinic polymorphs (Hughes
et al., 1989, 1990; Bauer & Klee, 1993).
An important part of the global phosphorus cycle is the
subduction of oceanic and continental lithospheric material along convergent plate boundaries. To understand the
residence and transport of P in the Earth’s deep interior,
the role of both phosphate and silicate phases has to be
considered. This is because under high P and T silicates
are capable of accommodating significant amounts of P
and, thus, become important subsolidus P carriers limiting
the stability fields of phosphate phases.
This study was undertaken to understand better the
crust^mantle recycling of P and halogens in subduction
zones and their storage and transport in the upper mantle
and Transition Zone. It involved (1) the determination of
the P^T stability of hydroxyl apatite and its high-T^highP breakdown products in bulk compositions typical for
subducting oceanic crust, (2) a study of the compositional
evolution of apatite and its breakdown products with P
and T, and (3) a study of P partitioning between phosphates and silicates under subsolidus and near-solidus conditions to assess the role of silicates as P storage sites in the
Earth’s mantle.
The geological significance of apatite
Apatite is the tenth most abundant mineral on Earth and a
widespread phase in almost all igneous and metamorphic
rocks. It carries an essential portion of the bulk P content
of 0·1^0·3 wt % P2O5 for continental crust (e.g. Weaver &
Tarney, 1984; Rudnick & Fountain, 1995; Rudnick & Gao,
2004; Kemp & Hawkesworth, 2004) and of 0·05^
0·25 wt % P2O5 for typical mid-ocean ridge basalts
(MORB; see below). In magmatic rocks originating from
deep-seated lithospheric and asthenospheric mantle
sources, such as lamproites, lamprophyres and kimberlites,
apatite is present as an abundant groundmass phase. This
may lead to bulk P2O5 contents 41·0 wt % with values as
high as 3·0 wt %. In many of these rocks P2O5 exceeds
Na2O (Mitchell & Bergman, 1991; Mitchell, 1995). Apart
from its importance as a P repository, apatite can also
accommodate significant amounts of rare earth elements
(REE) and large ion lithophile elements (LILE), sulfur,
carbon, and halogens (e.g. Edgar, 1989; Rnsbo, 1989;
Santos & Clayton, 1995; Parat & Holtz, 2004) and critically
influence the trace element evolution of magmas (e.g.
O’Reilly & Griffin, 1988, 2000; Piccoli & Candela, 2002).
NUMBER 11
NOVEMBER 2009
Apatite is absent from primitive mantle rocks because of
their low P contents of 90^100 ppm (McDonough & Frey,
1989; McDonough, 1990; McDonough & Sun, 1995; Palme
& O’Neill, 2004; Pearson et al., 2004; Workman & Hart,
2005). Metasomatically altered mantle rocks, however,
contain apatite as a common, albeit minor, constituent.
Apatite has also been described as inclusions in diamonds
(Lang & Walmsley, 1983; Guthrie et al., 1991). In many
instances phlogopite and/or Ca-amphibole coexisting with
apatite testifies to metasomatism by H2O-rich fluids or
melts (e.g. Dawson & Smith, 1977; Wass et al., 1980;
O’Reilly & Griffin, 1988, 2000; Ionov et al., 1997;
Morishita et al., 2003). According to Smith (1981) apatite is
the principal mantle reservoir for P and halogens.
O’Reilly & Griffin (2000) have also suggested that U^Thrich apatite can critically influence the isotopic evolution
and heat flow in metasomatically altered mantle segments.
The role of silicates as phosphorus
carriers in the crust and upper mantle
In silicates P commonly substitutes for silicon on the crystallographic T-sites with a strong preference for isolated
SiO4 tetrahedra in orthosilicates compared with the
linked tetrahedra in chain silicates (see Koritnig, 1965;
O’Neill & Mallmann, 2007). This behaviour is highly unusual for an incompatible trace element and makes garnet
and the Mg2SiO4 polymorphs potential storage sites for P.
P can be incorporated into garnet on a wt % level through
a coupled substitution (Mg,Ca)2þ þ Si4þ ¼ Naþ þ P5þ,
which is favoured at high pressures (e.g. Thompson, 1975;
Hermann & Spandler, 2007) and which would lead to an
Na-phosphate Na3Al2P3O12 with garnet structure (Thilo,
1941). In fact, Brunet et al. (2006) synthesized a complete
range of Mg3Al2Si3O12^Na3Al2P3O12 solid solutions in
the P^T range 15^17 GPa and 1200^16008C. The presence
in natural high-P rocks of garnet with P contents of up to
2000 ppm typically coupled with high Na contents is consistent with the experimental results (e.g. Bishop et al.,
1978; Schertl et al., 1991; Haggerty et al., 1994; Brunet &
Lecocq, 1999; Ye et al., 2000). P-rich garnets, however, are
not restricted to high-P environments. Up to 1·2 wt %
P2O5 have been reported from pegmatitic garnets by
Breiter et al. (2005), who proposed vacancy-producing substitutions of the type R2þ þ 2 Si4þ ¼ œ þ 2 P5þ to explain
the presence of P in virtually Na-free garnets.
P concentrations reported for olivine from common
mantle peridotites and basalts range from 550 to
1750 ppm, with most values 1000 ppm (Reid et al., 1975;
Bishop et al., 1978; Brunet & Chazot, 2001; Milman-Barris
et al., 2008). Exchange mechanisms proposed for P incorporation into olivine are 2 P þ œ(M1,2) ¼ 2 Si þ (Mg,
Fe)(M1,2) or 4 P þ œ ¼ 5 Si (Goodrich, 1984; Tropper
et al., 2004). P2O5 concentrations of up to several wt %,
rarely found in olivine from low-fO2 environments
(Buseck & Clark, 1984; Goodrich, 1984; Agrell et al., 1998;
2044
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
Tropper et al., 2004), are explained by disequilibrium
olivine^melt partitioning. P contents in ortho- and clinopyroxenes from mantle peridotites and eclogites are even
lower than those in garnet and olivine and range from
below the detection limit to 260 ppm (Reid et al., 1975;
Bishop et al., 1978; Brunet & Chazot, 2001). This is consistent with the finding by Koritnig (1965) that P in silicates
is inversely correlated with the degree of SiO4 tetrahedral
polymerization.
Previous experimental results on apatite
stability and phase relations
Most of the experimental studies of apatite stability have
investigated its near-solidus behaviour and its ability to
control the P, halogen, REE and LILE contents of the
coexisting partial melts (e.g. Watson, 1979, 1980; Baker &
Wyllie, 1992), or the partitioning of halogens between apatite and coexisting fluids or solids (e.g. Korshinsky, 1981;
Zhu & Sverjensky,1991,1992; Brenan,1993) under P^T conditions corresponding to the crust or the uppermost
mantle. In experiments under upper mantle P^T conditions Murayama et al. (1986) found that pure OH- and Fapatite breaks down to a previously unknown polymorph
of Ca3(PO4)2 and that the breakdown reaction has a negative slope. Sugiyama & Tokonami (1987) showed that this
new polymorph is isostructural with Ba3(PO4)2 (Roux
et al., 1978) and named the phase g-Ca3(PO4)2 by analogy
with the already known a’-, a-, and b-Ca3(PO4)2 polymorphs (Xie et al., 2003, and references therein). gCa3(PO4)2 was subsequently found as a shock-produced
natural phase in the Suizhou chondrite and named tuite
(Xie et al., 2002). A further phosphate with a Raman spectrum very similar to that of tuite had previously been
described by Chen et al. (1995) from the Sixiangkou chondrite. Because of its high Cl content, however, this phase
was identified as a high-P polymorph of Cl-apatite. No
experiments have been conducted so far to study the P^T
stability of apatite in representative peridotitic or basaltic
bulk compositions and no data on subsolidus P partitioning between phosphates and silicates under the P^Tconditions of the Earth’s mantle and Transition Zone have been
available. Likewise, nothing is known about the residence
of P in the lower mantle, although recently Brunet et al.
(2007) showed that six-coordinated P can be present in
stishovite, with the implication that P may be stored in
lower mantle silicates.
Experimental and analytical techniques
Two starting materials were used in the present study: (I)
an average MORB as reported by Melson et al. (1976) and
used by Yasuda et al. (1994); (II) a model Mg-basalt in the
system SiO2^TiO2^Al2O3^MgO^CaO^Na2O consisting
of a mixture of 30% garnet (prp90grs10) þ 60% clinopyroxene (di80jad10en10) þ 9% SiO2 þ1% TiO2 (Table 1).
This bulk composition allows all important exchange
reactions in an eclogitic assemblage to proceed but avoids
experimental and analytical problems associated with the
use of Fe-bearing starting materials. Both starting materials were prepared by mixing high purity (99·9%) and
finely ground SiO2, TiO2, MgO, CaCO3 and Na2CO3 in
ethanol for 20 min and stepwise decarbonation by heating
to 8008C with intermittent checks of the loss on ignition.
For higher reactivity, aluminum was added as g-Al2O3
after decarbonation. In a final step, 3 wt % synthetic OHapatite was added with homogenization for another
20 min. Bulk composition I was additionally doped with
10 trace elements (Nb, Ta, Zr, Y, Ba, Sr, Rb, Ce, Nd, Lu)
with concentrations between 250 and 350 ppm. The starting materials were stored at 2008C for at least 72 h and
then arc-welded into 2·0 or 1·6 mm outer diameter (o.d.)
Pt100 capsules. Two capsules were stacked in an individual
assembly and the length of a capsule did not exceed
1·7 mm for 2·0 mm o.d. capsules and 1·3 mm for 1·6 mm
o.d. capsules. An inner graphite liner was used for Febearing starting materials to minimize Fe loss and to provide an upper limit for fO2 (see Holloway et al., 1992;
Lesher et al., 2003; Konzett et al., 2008).
Experiments (Table 2) were performed with 500 t and
1000 t multi-anvil presses at the Institute of Mineralogy
and Petrology, University of Innsbruck (UI) and the
Bavarian
Research
Institute
of
Experimental
Geochemistry and Geophysics, University of Bayreuth
(BGI), respectively. The 500 t Walker-type multi-anvil
press operating at UI is a copy of the design used by the
BGI with assembly sizes and materials identical to those
employed at BGI. The multi-anvil press at the UI was calibrated to 10 GPa using the following phase transitions:
quartz^coesite (Bose & Ganguly, 1995), garnet^perovskite
in CaGeO3 (Susaki et al., 1985) and coesite^stishovite
(Zhang et al., 1996) with excellent agreement between UI
and BGI. Experimental and calibration conditions are
similar to those described by Rubie et al. (1993) and
Keppler & Frost (2005).
Phase compositions were analyzed with a JEOL 8100
superprobe using 10 nA beam current and 15 kV acceleration voltage. Natural minerals and synthetic oxides were
used for standardization in combination with the PRZ correction procedure. Measurement times were 20 s on peaks
and 10 s on backgrounds of the X-ray lines. For garnet
analyses, counting times for P, Na and Ti were increased
to 50 s/25 s (Table 3). Clinopyroxene was analyzed using
two sets of analytical conditions: in a first step the major
element composition was determined using 10 nA beam
current and counting times of 20 s on peaks and 10 s on
backgrounds of the X-ray lines. To avoid damage-induced
cation deficiencies a rastered beam was used. The accuracy
of the total cation sums of the pyroxenes was controlled
by analyzing a diospide standard (USNM 117733) under
identical conditions (see Konzett et al., 2008). In a second
2045
JOURNAL OF PETROLOGY
VOLUME 50
Table 1: Compositions of the starting materials
Bulk composition I
Bulk composition II
Average
MORB
model Mg-basalt
MORB
SiO2
48·3
53·8
TiO2
1·7
1·3
50·7
1·5
Al2O3
15·3
9·0
15·6
Fe2O3
1·0
—
FeO
8·3
—
MnO
0·2
—
MgO
8·2
18·3
7·7
CaO
13·0
15·0
11·4
Na2O
2·7
1·3
2·7
K2O
0·2
P2O5
1·2
1·2
H2O
P
0·1
0·1
100·0
100·0
—
1·0
8·9
—
NUMBER 11
NOVEMBER 2009
Quenched melt or fluid may be present as a mixture of
extremely fine-grained and needle-like quench phases dispersed between the solid phases or as a layer along the
interface between the graphite liner and the experimental
charge towards the hotter end of the capsule (Fig. 1e; see
Lesher & Walker, 1988). In bulk composition II the
quench is sufficiently abundant and homogeneous to be
analyzed with a high degree of reproducibility. Garnet
and clinopyroxene frequently show some compositional
variation in P, Ti and Fe/Mg as a result of incomplete
equilibration and/or partial Fe loss, especially in runs conducted at 510008C. When melt is present, Fe loss may be
severe despite the use of a graphite liner.
Phase relations
0·2
0·1
—
99·8
Bulk compositions I and II are doped with 3% synthetic
OH-apatite.
According to Melson et al. (1976).
analytical session, P was analyzed using the conditions
summarized in Table 3. The accuracy of the P analyses
was ensured by intermittent measurements of P-doped
glasses kindly provided by F. Brunet (see Brunet &
Chazot, 2001). In addition to the electron microprobe analysis (EMPA), polymorphs of SiO2 and TiO2 and the phosphate phases were identified using laser-Raman
spectroscopy.
R E S U LT S
Textures and chemical homogeneity of the
run products
All starting materials readily recrystallized to form euhedral to subhedral mineral grains. Clinopyroxene and the
SiO2 polymorphs usually form the largest grains with
20^100 mm diameter (Fig. 1). At P410 GPa, where
garnet is the dominant phase, the clinopyroxene grain
size can decrease to 510 mm. This often allowed only a few
reliable measurements to be made without beam overlap
with adjacent garnet. Compared with clinopyroxene,
garnet tends to form smaller grains, 10^20 mm in size,
rarely reaching 50 mm in runs conducted at the highest
temperatures. Apatite and tuite appear as rounded to elongated grains with an irregular grain shape and a size of
up to 50 mm. In many samples both apatite and tuite show
a strongly poikiloblastic texture (Fig. 1b and c) which
makes electron microprobe analyses of these phases difficult. At P410 GPa, tuite and the TiO2 polymorphs are
typically present as needles 53^5 mm in diameter and,
hence, too small for electron microprobe analysis (Fig. 1a).
Bulk composition I crystallized to form garnet þ
omphacite þ coesite/stishovite þ rutile/TiO2-II þ apatite/
tuite in the P^Trange 7^11 GPa and 950^12008C, thus representing a phosphate-bearing eclogite. Towards higher
P and T, clinopyroxene, TiO2-II and the phosphate phases
disappear as a result of partitioning of Na, P and Ti into
garnet (Fig. 2a, Table 2). This process forms a stishovitebearing garnetite at 13^15 GPa and 1350^16008C. The
modal amount of garnet present in this P^T range cannot
accommodate the entire bulk P budget as is obvious from
the P2O5 contents of garnet. The remaining P could be
stored in trace amounts of tuite not detected in the central
section through the capsule or undetected fluid residing
along grain boundaries of the graphite liner. Textural evidence for fluid formed by apatite breakdown and/or moisture adsorbed onto the starting material is provided by
small amounts of quench phases (see above). It is only in
run B07-16 at 15 GPa and 18008C that substantial amounts
of quenched melt are present, coexisting with garnet and
stishovite.
Bulk composition II crystallized to form similar assemblages to composition I, except at P 4 GPa where additional small amounts of enstatite and magnesite are
present (Fig. 1b and Table 2). The latter formed from trace
amounts of CO2 inherited from the starting materials.
Clinopyroxene, TiO2 and the phosphate phases are stable
to slightly higher P and T of 13 GPa and 13508C compared
with bulk composition I (Fig. 2b and Table 2). All runs
using bulk composition II show small amounts of
quenched fluid or melt.
Phase compositions
Averaged and representative compositions of phases from
all experiments are given as Supplementary Data at
http://www.petrology.oxfordjournals.org/.
Apatite and tuite
In bulk composition I apatite is present in one experiment
at 7 GPa and 9508C very close to its upper P stability
limit where it is the major host for Sr and LREE
2046
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
Table 2: Summary of experimental run conditions and products
Run
Bulk
Assembly
no.
P
T
(GPa)
(8C)
Position
Duration
Phases observed
MA36
I
18/11
7·0
950
lower
363 h 15 min
grt þ cpx þ coe þ rt þ ap
B08-7
I
18/11
7·0
950
lower
190 h 47 min
grt þ cpx þ coe þ rt þ ap
MA37
I
18/11
7·0
1100
upper
262 h 20 min
grt þ cpx þ coe þ rt þ Q
MA41
I
18/11
7·5
950
lower
260 h 15 min
grt þ cpx þ coe þ TiO2-II þ tu þ Q
B06-9
I
18/11
8·0
950
upper
313 h 20 min
grt þ cpx þ coe þ TiO2-II þ tu
B05-17
I
18/11
8·0
1050
upper
186 h 07 min
grt þ cpx þ coe þ TiO2-II þ tu (þ mag)
B07-15
I
18/11
8·0
1125
upper
101 h 04 min
B05-10
I
18/11
8·0
1200
upper
94 h 50 min
grt þ cpx þ coe þ rt þ tu (þ mag)
grt þ cpx þ coe þ TiO2-II þ Q
B05-19
I
18/11
11·0
1100
upper
168 h 25 min
grt þ cpx þ stish þ TiO2-II þ tu
B05-3
I
18/11
11·0
1200
upper
72 h 45 min
grt þ cpx þ stish þ TiO2-II þ tu
B05-1
I
18/11
11·0
1350
upper
24 h 10 min
grt þ cpx þ stish
B05-5
I
14/8
13·0
1350
upper
23 h 10 min
grt þ stish
B05-7
I
14/8
15·0
1450
upper
14 h 00 min
grt þ stish
B07-17
I
14/8
15·0
1600
upper
05 h 00 min
grt þ stish
B07-16
I
14/8
15·0
1800
upper
04 h 22 min
grt þ stish þ Q
JKI-67
II
PC
3·0
850
341 h 57 min
grt þ cpx þ opx þ coe þ rt þ ap þ Q
JKI-91
II
PC
3·0
1000
508 h 06 min
grt þ cpx þ rt þ Q
MA32
II
18/11
4·0
900
lower
303 h 00 min
grt þ cpx þ opx þ coe þ rt þ ap þ Q
MA30
II
18/11
4·0
950
upper
334 h 45 min
grt þ cpx þ opx þ rt þ Q
MA23
II
18/11
5·0
950
lower
285 h 12 min
grt þ cpx þ coe þ rt þ ap þ Q (þ mag)
MA09
II
18/11
6·0
950
upper
185 h 51 min
grt þ cpx þ coe þ rt þ ap þ Q (þ mag)
MA51
II
18/11
6·0
1100
lower
145 h 55 min
grt þ cpx þ coe þ rt þ Q
MA26
II
18/11
7·0
950
lower
217 h 04 min
grt þ cpx þ coe þ rt þ ap þ Q (þ mag)
B08-7
II
18/11
7·0
950
upper
190 h 47 min
grt þ cpx þ coe þ rt þ ap þ Q (þ mag)
MA37
II
18/11
7·0
1100
lower
262 h 10 min
grt þ cpx þ coe þ rt þ Q
MA41
II
18/11
7·5
950
upper
236 h 15 min
grt þ cpx þ coe þ TiO2-II þ tu þ Q (þ mag)
B06-9
II
18/11
8·0
950
lower
313 h 20 min
grt þ cpx þ coe þ TiO2-II þ tu þ Q (þ mag)
B07-15
II
18/11
8·0
1125
lower
101 h 04 min
grt þ cpx þ coe þ rt þ tu þ Q (þ mag)
MA16
II
18/11
10·0
1200
upper
120 h 20 min
grt þ cpx þ stish þ TiO2-II þ tu þ Q (þ mag)
B05-3
II
14/8
11·0
1200
lower
72 h 54 min
B05-1
II
14/8
11·0
1350
lower
24 h 10 min
grt þ cpx þ stish þ TiO2-II þ tu þ Q (þ mag)
B05-5
II
14/8
13·0
1350
lower
23 h 10 min
grt þ cpx þ stish þ TiO2-II þ tu þ Q (þ mag)
B05-7
II
14/8
15·0
1450
lower
14 h 00 min
grt þ cpx þ stish þ Q
grt þ cpx þ stish þ TiO2-II þ tu þ Q þ (þ mag)
grt, garnet; cpx, clinopyroxene; rt, rutile; TiO2-II, TiO2 with a-PbO2 structure; coe, coesite; stish, stishovite; ap, apatite;
tu, tuite; mag, magnesite; Q, quenched fluid or melt; PC, piston cylinder. Bulk compositions are given in Table 1.
Position of the charge in the MA assembly with respect to the thermocouple (TC): upper, close to TC; lower,
far from TC.
(0·26 0·03 wt % SrO, 1·03 0·11 Ce2O3 and
1·12 0·06 wt % Nd2O3). In addition, significant MgO,
FeO and SiO2 and Cl are present, the latter most likely
inherited from the graphite used for the graphite liner.
Spectrometer scans yielded no evidence for Rb, Ba, Lu
and Y at concentration levels measurable by EMPA. Tuite
from bulk composition I has distinctly higher Sr, Ce, Nd
and Na but does not contain measurable Mg (Fig. 3b).
In bulk composition II the only minor element measurable
in apatite is Mg. Its concentration increases with increasing P from 1·0 wt % MgO at 3 GPa and 8508C to 1·8 wt %
MgO at 7 GPa and 9508C, which is equivalent to
4·4 mol % Mg5(PO4)3(OH) solid solution (Fig. 3a). This
increase in Mg with increasing P is consistent with the
smaller ionic radius of Mg2þ compared with Ca2þ, irrespective of the coordination number (Shannon, 1976).
2047
JOURNAL OF PETROLOGY
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Table 3: Analytical conditions for EMPA of garnet and
clinopyroxene
Garnet
Clinopyroxene
Acceleration voltage
15 kV
15 kV
Beam current
10 nA
150 nA
Analytical conditions for P
X-ray line/crystal
P-Ka/PETH
P-Ka/PETH
Counting time
50 s/25 s
100 s/50 s
P standard
natural F-apatite
natural F-apatite
Detection limit (ppm P)
70
15
Typical 2s error
12% for 1000 ppm P
8% for 200 ppm P
Standards and counting times (peak/background) for remaining elements
NUMBER 11
NOVEMBER 2009
compositions (e.g. Okamoto & Maruyama, 2004;
Spandler et al., 2007; Konzett et al., 2008). In addition, up
to 6 mol % Ca-Eskola pyroxene solid solution are present
between 11008C and 12008C. In bulk composition II, clinopyroxene is a diopside^jadeite solid solution with 75^
82 mol % diopsidess and 10^15 mol % jadeitess. There is
very little compositional variation across the P^T region
investigated and no significant Ca-Eskola solid solution is
present. The P concentration in clinopyroxene ranges
between 129 14 and 258 25 ppm P and slightly increases
with P and T (Fig. 5). Clinopyroxenes 50 mm may show
variations in their P contents that are outside the 2s error
of a single analysis (Table 3). These variations, however,
do not represent systematic core-to-rim zoning trends. No
relation between major element composition and P concentrations could be observed (compare runs MA36 and
MA26).
Si
pyrope (20/10)
diopsidey (20/10)
Ti
rutile (50/25)
rutile (20/10)
Al
pyrope (20/10)
corundum (20/10)
Accessory phases
Mg
diopside (20/10)
diopside (20/10)
Ca
pyrope (20/10)
diopside (20/10)
Fe
pyrope (20/10)
pyrope (20/10)
Mn
tephroite (20/10)
tephroite (20/10)
Na
jadeite (50/25)
jadeite (20/10)
Stishovite always contains significant Al2O3; it ranges
between 0·3 and 0·9 wt % and increases with increasing
T (Liu et al., 2006b; Litasov et al., 2007). Stishovite from
bulk composition I also contains Fe3þ with a molar Al:Fe
ratio of 1:1. The presence in stishovite of Al and Fe3þ is consistent with hydrogen incorporated through a substitution
(Al3þ, Fe3þ) þ Hþ ¼ Si4þ (see Pawley et al., 1993; Litasov
et al., 2007, and references therein). No P could be detected
in stishovite using the analytical procedure applied to clinopyroxene. All TiO2 polymorphs contain significant
Al2O3 and CaO and those from bulk composition I are
also major hosts for Nb2O5, Ta2O5, and ZrO2. For a discussion of minor element incorporation into the TiO2 polymorphs the reader is referred to Vlassopoulos et al. (1993)
and Bromiley & Hilairet (2005). Quenched melt or fluid
from bulk composition II is siliceous with 37^41wt %
SiO2 at P 6 GPa and has very high P2O5 contents of
5·9^7·6 wt %. At 15 GPa and 14508C the melt is carbonatitic with 2 wt % SiO2 and is again strongly enriched in
P2O5 (5·1wt %).
USNM standard 143968.
yUSNM standard 117733 (see Jarosevich et al., 1980).
By analogy with bulk composition I, Mg in tuite is always
below the detection limit, yielding stoichiometric
Ca3(PO4)2 in all experiments.
Garnet
In bulk composition I garnet forms a complex pyrope^
almandine^grossular solid solution with significant
amounts of P, Na and Ti. In the presence of apatite or
tuite P increases with increasing P and T from 0·2 wt %
P2O5 at 7 GPa and 9508C to 0·6 wt % P2O5 at 11 GPa
and 12008C.
A maximum of 0·78 0·06 wt % P2O5 in the garnet is
reached at P 13 GPa (Fig. 4a). In these experiments tuite
could not be detected in the central section through the
capsule. In bulk composition II P2O5 remains constant at
0·2^0·3 wt % to P 10 GPa and then increases to 0·6 wt %
at 13 GPa and 13508C (Fig. 4b). Possible mechanisms
responsible for P^Na^Ti incorporation into garnet are discussed below.
Clinopyroxene
In bulk composition I clinopyroxene forms a diopside^
jadeite solid solution with minor Ca-Tschermak’s pyroxene
and enstatite solid solution. Jadeitess increases from 43 mol
% at 7 GPa to 67 mol % at 11 GPa, which is in accordance
with earlier studies that used comparable bulk
Raman spectra of tuite and apatite
Raman spectra of tuite show a very intense band at 976^
979 cm^1 with less intense bands at 1094^1098 cm^1, 1001^
1005 cm^1, 639^643 cm^1, 577^579 cm^1 and 411^414 cm^1.
This is in very good agreement with band positions
reported by Xie et al. (2002, 2003) for natural tuite and
also for synthetic g-Ca3(PO4)2. Apart from these bands
found in 22 spectra from seven experiments, weak bands
at 1095^1097 cm^1, 520^525 cm^1, 192^194 cm^1, 176^
177 cm^1, 150^153 cm^1 and in the range 414 and 192 cm^1
may also be present (Fig. 6a). These are thought to originate from undetected carbonate or silicate inclusions in
tuite.
The Raman spectrum of apatite (e.g. Nelson &
Williamson, 1982) can be easily distinguished from that of
2048
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
Fig. 1. Back-scattered electron photomicrographs of run products from the high-P experiments: (a) bulk composition II; run MA23 at 5 GPa
and 9508C; (b) bulk composition II, run MA41at 7·5 GPa and 9508C; (c) bulk composition I, run MA36 at 7 GPa and 9508C; (d) bulk composition I, run MA41at 7·5 GPa and 9508C; (e) bulk composition II, run MA30 at 4 GPa and 9508C. Abbreviations as in Fig. 2.
tuite based on the shift of the major v1 PO4-stretching
mode from 975 to 962 cm^1 (Fig. 6a and b). In apatite
from bulk composition I a small band at 3508 cm^1 is present in addition to the main band at 3570 cm^1. This is
consistent with the presence of small amounts of Cl (see
Elliott, 1994). In spite of the presence of magnesite, the
Raman spectra of apatite from bulk composition II do not
provide conclusive evidence for CO3 groups.
2049
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DISCUSSION
The residence of phosphorus in basaltic
rocks
All volcanic and plutonic rocks of basic and intermediate
composition contain small amounts of P usually of the
order of51wt % P2O5 (LeMaitre,1976). MORBs typically
have 0·05^0·25 wt % P2O5 (e.g. LeMaitre, 1976;
Wilkinson & LeMaitre, 1987; Hart et al., 1999; Klein,
2004; Workman & Hart, 2005). Basalts and andesites from
island arcs and active continental margins show a wider
range in P concentrations from 0·1 to 0·7 wt % P2O5
(e.g. LeMaitre, 1976; Wilkinson & LeMaitre, 1987; Condie,
1993; Plank & Langmuir, 1998; Kelemen et al., 2004).
Even higher concentrations of 0·5^1·6 wt % P2O5 are
found in continental alkali-rich and feldspathoid-bearing
basalts (LeMaitre, 1976; Wilkinson & LeMaitre, 1987;
Farmer, 2004). P increases with increasing degree of differentiation together with Fe and Ti, which may lead to bulk
P2O5 concentrations of several wt % (e.g. Leeman et al.,
NUMBER 11
NOVEMBER 2009
1976; Hill, 1988; McLelland et al., 1994; Costa & Caby,
2001; Sarapa«a« et al., 2005; White, 2007).
Major hosts for P are olivine, clinopyroxene and plagioclase. The P concentration in these phases is a complex
function of P, T, oxygen fugacity and SiO2 content of the
coexisting melt (Libourel et al., 1994; Bindeman et al.,
1998; Milman-Barris et al., 2008). Because of the high
P-saturation concentration in basaltic melts, apatite is unlikely to be a phenocryst phase stable near the liquidus at
the time of extrusion (Watson, 1979). Nevertheless,
trace amounts of apatite are very common in basalts or
gabbros (e.g. Anderson & Greenland, 1969; Neumann
et al., 2000; Coogan et al., 2001; Meurer & Natland, 2001;
Thy, 2003; Kaczmarek et al., 2008) and may form by a
local build-up of P sufficient to reach P saturation (Green
& Watson, 1982). In rare instances rocks with 30^40 vol.
% apatite coexisting with FeTi-oxides may form
from immiscible liquids that separated from strongly differentiated anorthositic^mangeritic magmas (Philpotts,
1967).
Fig. 2. Schematic P^T diagrams summarizing experimental results. (a) Results for bulk composition I (Table 1); phases present in the experimental charges are represented by black or grey sectors within the run symbol; phases not detected are denoted by white sectors (see inset
upper left). Abbreviations: SiO2, coesite or stishovite; cpx, clinopyroxene; TiO2, rutile or TiO2-II; ap, apatite; Q, quenched fluid or melt; grt,
garnet; tu, tuite; M86, apatite^tuite reaction for pure hydroxyl-apatite according to Murayama et al. (1986); ACMA, average current mantle
adiabat. (b) Results for bulk composition II; symbols and abbreviations as in (a), and mag, magnesite; opx, orthopyroxene.
2050
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
The role of apatite and tuite as phosphorus
and LILE^REE carriers in subducted
basalts
Fig. 2. Continued
During high-P metamorphism of MORB-type rocks P is
redistributed from olivine and plagioclase to clinopyroxene
and garnet. At P52 GPa both clinopyroxene and garnet
are unable to store significant amounts of P. Therefore,
this redistribution of P results in the formation of apatite
in excess of that inherited from the basalt precursor.
Monazite and xenotime are very rare in basalts (Lovering
et al., 1973; Herd et al., 2004) as a result of bulk compositional constraints and, hence, do not play any significant
role as P carriers in these rocks.
By the time the basalt or gabbro reaches eclogite-facies
conditions, a major portion of the bulk P content of the
rock will be stored in apatite whereby 0·5% apatite is sufficient to store the entire MORB P budget of 0·2 wt %
P2O5. In fact, apatite is an extremely common accessory
phase in eclogites and has been reported from worldwide
high-P (HP) and ultrahigh-P (UHP) localities of both
crustal and mantle origin (e.g. Brueckner et al., 1998;
Svensen et al., 2001; Heaman et al., 2002; Zack et al., 2002;
Spandler et al., 2003; Baldwin et al., 2004; Zhang et al.,
2005; Srensen, 2006; Aulbach et al., 2007; Sun et al., 2007).
At P43 GPa garnet starts to incorporate significant
amounts of P (see above), which leads to a decrease in the
modal amount of apatite.
Fig. 3. Compositional variation with P and T of apatite and tuite. (a) Variation of the averaged MgO content in apatite for bulk
composition II; numbers next to symbols are experimental temperatures; at 9508C two data points are plotted to compare the influence
of the capsule position (u, upper; l, lower; see section on experimental and analytical conditions) on phase compositions. (b) Variation of averaged minor element concentrations in apatite and tuite in bulk composition I across the ap^tu reaction and in an isobaric section at 8
GPa; bars next to element abbreviations give 2s errors for these elements in a single EMP analysis under analytical conditions as outlined in
the text.
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Because of the ability of garnet to incorporate significant
P under UHP conditions, the question arises whether or
not apatite will reach the P(^T) conditions of the apatite^
tuite reaction in a subducted MORB. To assess this problem the modal composition of run B08-7 (bulk composition I; 7 GPa and 9508C) was determined and the P
storage capacity of garnet and clinopyroxene was calculated based on their averaged P2O5 concentrations (see
Supplementary Data). The results (Table 4) show that
NUMBER 11
NOVEMBER 2009
only 0·3% apatite is sufficient to supply all the P for
garnet and clinopyroxene. If it is assumed that the bulk
P2O5 content of 0·2 wt % in an average MORB is stored
in apatite then only 60% of the apatite will have broken
down by the time the P^T conditions of the apatite^tuite
breakdown are reached. In rocks with a higher bulk P content the persistence of apatite is even more likely. It has to
be emphasized, however, that this calculation is valid
only for subsolidus conditions because of the rapid
Fig. 4. Variation with P and T in the averaged P2O5 content of garnet (a) from bulk composition I and (b) from bulk composition II; numbers
next to symbols are temperatures (8C); filled symbols represent P contents of garnets coexisting with apatite or tuite; dashed line gives approximate P of the ap^tu reaction.
2052
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
consumption of apatite in basic to intermediate melts close
to the solidus (Watson, 1979). A further potential subsolidus
P host not considered in the calculation is the fluid phase.
Its role is difficult to assess because no data on P solubility
in high-P fluids have been available so far. Run B08-7
(bulk composition II), however, provides evidence for limited solubility of P. Although needle-like quench phases
show that a fluid phase was present during the experiment,
abundant apatite crystals are still stable (Table 2). This
indicates that small amounts of fluid are not able to
induce the breakdown of significant amounts of apatite.
The results of this study show that apatite has a much
more restricted stability range than that indicated by
Murayama et al. (1986). Hydroxyl apatite is not stable
above 7·5 GPa at 9508C, which corresponds to a depth of
200 km. This is because silicate phases or phase components are involved in the tuite-forming reaction (see
Brunet et al., 1999). During apatite breakdown both water
and halogens are liberated (see below) whereas LILE and
REE are retained and even concentrated in tuite, the
latter taking over the role as a major P, LILE and REE
carrier at depths 4200 km (see Fig. 3; see also Murayama
et al., 1986; Sugiyama & Tokonami, 1987).
The stoichiometry of the apatite^tuite reaction cannot
be identified because of the absence of breakdown products
in addition to those present (grt þ cpx þTiO2 þ SiO2)
and the lack of significant compositional changes of
garnet and clinopyroxene owing to the small amount of
apatite in the starting material. If apatite þ tuite þ pyrope þ
grossular þ diopside þ enstatite þ coesite þ H2O are considered as phases or phase components then the possible
tuite-forming reactions are
12 apatite þ pyropess þ 6 coesite ¼
18 tuite þ grossularss þ 3 diopsidess þ 6 H2 O
12 apatite þ 2 pyropess þ6 coesite ¼
18 tuite þ 2 grossularss þ3 enstatitess þ 6H2 O
4 apatite þ enstatitess þ2 coesite ¼
Fig. 5. Averaged P contents in clinoproxene from bulk compositions I
and II as a function of P and T.
6 tuite þ 2 diopsidess þ2 H2 O:
Fig. 6. Unpolarized Raman spectra (a) for tuite and (b) for apatite from bulk compositions I and II; bold numbers give positions of Raman
bands as observed in this study, numbers in parentheses give band positions observed by Xie et al. (2003) for tuite and by Nelson &
Williamson (1982) for apatite.
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This study confirms the conjecture of Sugiyama &
Tokonami (1987) that both apatite and tuite are important
carriers of LILE and REE to depths in the upper mantle
and Transition Zone. A comparison of minor element variations in apatite and tuite in bulk composition I indicates
different incorporation mechanisms for the LREE in both
phases. In tuite, Ce and Nd are accompanied by significant Na. Their co-variation, including data from an
incompletely equilibrated tuite core (see Fig. 7a), which is
close to a 1:1 trendline, indicates a substitution
REE3þ þ Naþ ¼ 2 Ca2þ (Rnsbo, 1989; Pan & Fleet,
2002). By comparison, apatite, although showing significant Ce þ Nd contents, has very low Na contents, which
are inconsistent with a substitution REE3þ þ Naþ ¼
2 Ca2þ. Instead, significant Si contents point to a substitution REE3þ þ Si4þ ¼ Ca2þ þ P5þ (Fig. 7b) (Pan & Fleet,
2002). For tuite the potential role of Si in coupled substitutions cannot be assessed with confidence because almost
all tuites show tiny inclusions of silicate phases, which
make a reliable determination of Si impossible. In any
case the Na^REE systematics in tuite do not require a significant contribution of Si-involving REE exchange
mechanisms.
Table 4: P budget for run B08-7 (7 GPa and 9508C)
using bulk composition I
The role of apatite as halogen and water
carrier in subducted basalts
Mineral phases present
Modal amount (%)
grt
cpx
54
32
coe
rt
8
1
Average P2O5 concentration (wt %)
0·19
0·039
—
—
Amount of P2O5 stored (wt %)
0·103
0·012
—
—
6·0
—
—
ap
3
Fraction of typical MORB-type P2O5
budget of 0·2 wt % (%)
51·5
Amount of ap required to supply
0·12 wt % P2O5 (wt %)
0·28
Amount of ap required to store
0·2 wt % P2O5 (wt %)
0·47
Modal amounts calculated from bulk and averaged mineral
compositions (Table 1 and Electronic Appendix).
Apatite is an important potential subsolidus halogen reservoir during high-P metamorphism and crust-to-mantle
transport of basalts in subduction zones, but it does not
play any significant role as a water reservoir (Fig. 8).
Whereas apatite from felsic igneous rocks is mostly F-rich,
that from mafic rocks shows a wide range in OH^F^Cl
ratios (e.g. Wass et al., 1980; Exley & Smith, 1982; Ionov
et al., 1997; O’Reilly & Griffin, 1988, 2000; Piccoli &
Candela, 2002). Unlike silicates, halogen incorporation
into apatite is not subject to crystal chemical constraints
(Mg^Cl avoidance; e.g. Volvinger et al., 1985; Oberti et al.,
1993; Kullerud, 1995). Primitive MORB has a bulk Cl content of 20^50 ppm (Oppenheimer, 2004, and references
therein). To store this amount of Cl, 0·4^1·1wt % Cl in apatite is required if it is assumed that the typical bulk P2O5
concentration of 0·2 wt % is entirely locked up in apatite
(Fig. 8). The very few reports on halogen contents of eclogitic apatite show a wide range in Cl from 0·02 to 1·7 wt %
Fig. 7. Compositional variation of REE^Na^Si with P and T in apatite and tuite from bulk composition I: (a) Na vs (Ce þ Nd) for both apatite
(filled symbols) and tuite (open symbols); numbers next to symbols in inset lower right are P (GPa)/T (8C); (b) Si vs (Ce þ Nd) for apatite.
2054
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
(Svensen et al., 2001; Zack et al., 2002; Miller et al., 2007;
Glodny et al., 2008). This shows that apatite is indeed a
major host for Cl in subducted MORB. Apart from apatite,
the only solid phase stable to depths 100 km that can
accommodate Cl on a wt % level is halite. This phase is
occasionally found in eclogites as daughter crystals in
saline fluid inclusions (e.g. Philippot et al., 1998, and references therein; Svensen et al., 2001) and it may also be present in the matrix of eclogites. Unfortunately, the common
method of wet polishing of thin sections will dissolve
away any trace amounts of matrix halite. Hence, halite
may be more widespread in eclogites than is commonly
assumed based on evidence from fluid inclusions. A mechanism that may form halite during subduction is progressive
desiccation of saline fluid inclusions as a result of the
increasing OH solubility in clinopyroxene with increasing
P (see Katayama & Nakashima, 2003).
The role of phengite and nominally anhydrous phases as
Cl carriers in eclogites is difficult to assess because of a
lack of data. Svensen et al. (2001) have shown that even in
the presence of concentrated brines the Cl contents of
phengite are 0·03 wt %. It is only in very rare instances
that phengite with 40·1wt % Cl is reported from eclogites
(John & Schenk, 2003). Secondary ionization mass spectrometry (SIMS) analyses of clinopyroxenes from peridotites have yielded highly variable Cl contents of 20^
400 ppm (Scambelluri et al., 2004; Ottolini & Le Fe'vre,
2007). Whether Cl contents of the order of several hundred
Fig. 8. Cl and H2O storage capacity of 0·47 wt % (OH, Cl)-apatite
(amount of apatite required to store average MORB P budget of
0·2 wt % P2O5) as a function of mol % Cl-ap or wt % Cl in apatite
(upper scale); horizontal grey bar represents range of average Cl concentration in MORB reported by Oppenheimer (2004); vertical grey
bar gives Cl concentration in apatite required to store the corresponding average bulk Cl content.
ppm are incorporated into the pyroxene lattice or originate
from sub-microscopic fluid inclusions is not yet clear.
The subsolidus breakdown of apatite at a depth of
200 km is a barrier for the crust-to-mantle transport of
halogens in subducted basalts and will result in a fractionation of F and Cl. Whereas OH and F can at least in part
be redistributed into nominally anhydrous phases such as
omphacite (see Koch-Mu«ller et al., 2004; Katayama et al.,
2006) and/or hydrous silicate phases (e.g. phengite, topazOH, Mg-pumpellyite, lawsonite) Cl is most probably lost
to the fluid. This is because of its strong incompatibility in
hydrous silicates except for very Fe-rich bulk compositions.
Thus, apatite breakdown contributes to the Cl depletion
of subducted oceanic crust and the reflux of Cl to the surface (Straub & Layne, 2003).
Phosphorus in eclogitic silicate phases
In a plot of Na vs (P þ Ti) (Fig. 9a) for garnet from bulk
composition I, the data points are aligned very close to a
1:1 trendline in the P^T range 7^8 GPa and 950^12008C.
Fig. 9. Variation of the minor element composition of garnet from
bulk composition I with P and T: (a) Na vs (Ti þ P); (b) Na vs
(P þ Ti þ [6]Si). In legend, numbers next to symbols are P (GPa)/T
(8C).
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JOURNAL OF PETROLOGY
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Fig. 10. Variation of the minor and major element composition of garnet from bulk composition I: (a) with T from 950 to 12008C in an isobaric
section at 8 GPa; (b) and (c) for garnet from run B05-10 at 8 GPa and 12008C.
This indicates that Na is fully compensated by substitutions
[8]
Na[4]P[8]M2þ^1[4]Si^1 and [8]Na[6]Ti[8]M2þ^1[6]Al^1 as
observed in earlier studies (e.g. Thompson, 1975; Haggerty
et al., 1994; Brunet et al., 2006; Hermann & Spandler,
2007). At P48 GPa the data points increasingly deviate
from a 1:1 towards a 2:1 trend. This suggests that substantial Na not associated with P and Ti is incorporated into
garnet. The excess of Na compared with (P þ Ti)
(Fig. 9b) can be attributed to a substitution
[8]
Na[6]Si[8]M2þ^1[6]Al^1 (Ringwood & Major, 1971),
which introduces a majorite component. The unusually
low Na, Ti and P concentration in garnet from B07-16 is a
result of partitioning of these elements into coexisting
melt. In an 8 GPa isobaric series of experiments there is
little variation in P between 9508C and 11258C as a result
of buffering by tuite. By contrast, both Ti and Na strongly
increase with T. Figure 10a shows that P^Na^Ti variation
in garnet can be largely attributed to a combination of
NaPM2þ^1Si^1 and NaTiM2þ^1Al^1 between 950 and
11258C. It is only in the 12008C experiment (B05-10) that
data points start to significantly deviate from the 1:1 trend
in (Ti þ P) vs Na towards higher Ti þ P. The negative Ti^
Ca and Ti^Al correlations for garnet from B05-10
(Fig. 10b and c) are consistent with Ti incorporation
through an exchange [8]Ca[6]Ti[6]Al^2 (see Zhang et al.,
2003).
Garnet analyses from bulk composition II also show a
positive correlation between Na, P and Ti with Na and P
plotting on a 1:1 trendline at P510 GPa and T511258C
(Fig. 11a). This indicates that NaPM2þ^1Si^1 accounts for
most if not all Na and P in the garnet. Towards higher
P and T, Ti becomes increasingly important through
NaTiM2þ^1Al^1. In contrast to bulk composition I, however, there is no indication for [8]Na[6]Si[8]M2þ^1[6]Al^1.
This is because all garnet analyses, even those with up to
0·3 [6]Si a.p.f.u., plot on or above a 1:1 trendline in the Na
vs (P þ Ti) diagram (Fig. 11b). Data from experiments at
13 and 15 GPa show the strongest deviation towards high
(Ti þ P) and a strong negative correlation between [6]Si
and Ti. This would be consistent with Ti incorporation
through an exchange [6]M2þ[6]Ti[6]Si^2.
Averaged analyses of coexisting quenched melt or fluid
and garnet from runs at 4, 6 and 15 GPa in bulk composition II were used to calculate partition coefficients for P.
This is justified because of the compositional homogeneity
of the quench, which is comparable with that of the solid
phases. The resulting DPgrt^melt/fluid is 0·05 and 0·03 at 4
and 6 GPa. The strong increase of P in garnet with increasing P yields DPgrt^melt/fluid ¼ 0·11at 15 GPa. It is interesting
to note that the P2O5 content of the carbonatitic melt or
fluid is lower than that of the siliceous melts or fluids.
Compared with melt and garnet, the P storage capacity
of clinoproxene is negligible in the investigated bulk compositions across the entire pyroxene P stability interval.
Between 4 and 11 GPa, DPgrt^cpx shows little variation
between 5·9 and 7·5 when buffered by apatite or tuite. In
the absence of a detectable phosphate phase DPgrt^cpx may
increase to 10·4 as a result of the increasing P^(Na^Ti) solubility in garnet close to the upper P stability limit of
clinopyroxene.
Brunet et al. (2007) synthesized stishovite with 1wt %
P2O5 at 18 GPa and 16008C and showed that P in this
2056
KONZETT & FROST
HIGH P^T STABILITY OF OH-APATITE
P^Ti^(Na)-rich garnet (Haggerty et al., 1994; Ye et al.,
2000). Likewise, the DPgrt^cpx values 51 calculated for
coexisting garnet and clinopyroxene from mantle eclogites
reported by Bishop et al. (1978) are a strong indication
for disequilibrium element partitioning. Thus, the
comparison of Na^Ti^P contents of garnets from UHP
metapelites and metabasic rocks with experimental data
indicates that garnet rarely preserves its high-P^high-T
equilibrium composition. This has to be kept in mind
when thermobarometric data are retrieved from UHP
assemblages.
S U M M A RY A N D C O N C L U S I O N S
Fig. 11. Variation of the minor element composition of garnet from
bulk composition II with P and T: (a) Na vs P; (b) Na vs (P þ Ti).
Numbers next to symbols are P (GPa)/T (8C).
phase is present in six-fold coordination. This study, however, indicates that stishovite coexisting with garnet is unlikely to be a major storage site for P because of the
extremely strong preference of P for the garnet structure.
Hermann & Spandler (2007) pointed out the discrepancy between significant Na, Ti, and P contents of metapelitic garnets experimentally equilibrated at 2·5^4·5 GPa
and 600^10508C and the very low concentration of these
elements in the majority of natural HP and UHP rocks.
This was attributed to substantial re-equilibration during
the comparatively slow uplift and exhumation of continental crustal rocks. The same discrepancy can be observed
for the garnets of this study compared with those from
many natural UHP crustal and mantle eclogites (e.g.
Reid et al., 1976; Bishop et al., 1978; Nowlan et al., 2000;
Zhang et al., 2003; Katayama et al., 2006; Liu et al., 2006a).
A (partial) removal of P, Na and Ti from garnet therefore
seems to be a widespread phenomenon even in mantle
eclogites despite their extremely rapid ascent to the surface. In fact, evidence for this removal is occasionally provided by oriented apatite and rutile inclusions in garnet
that are interpreted as exsolution from an originally
(1) Apatite is an important storage site for LILE, REE and
halogens in subducted basaltic oceanic crust. In an average
MORB and a model Mg-basalt its stability is limited to
57·5 GPa at 9508C in a typical eclogite assemblage
garnet þ clinopyroxene þ SiO2 þ TiO2. This is 5 GPa
lower than the upper P stability limit of pure OH-apatite
and corresponds to a depth of 200 km. The high-P breakdown product of apatite is tuite [g-Ca3(PO4)2].
(2) Both apatite and tuite are stable in a wide range of
subduction zone T regimes but not along an average
mantle adiabat. This precludes apatite or tuite stability in
the asthenospheric mantle. Apatite may be stable in cold
continental lithosphere (40 mW/m2) but is restricted to P
5 4^5 GPa.
(3) Apatite breakdown is an important limit for Cl
transport in subduction zones and can contribute to the
Cl depletion of subducted crust. This is because with the
exception of (Na, K)Cl no subsolidus phases are available
in an average MORB that can accommodate significant
Cl at P47 GPa. In comparison, apatite breakdown has
much less impact on F and H2O storage and transport in
subduction zones because both H and F can be accommodated by hydrous phases and/or nominally anhydrous
minerals.
(4) Apatite shows a continuous and bulk compositiondependent increase in MgO with increasing P reaching
1·9 wt % at 7 GPa and 9508C equivalent to 4·4 mol %
Mg5(PO4)3(OH) solid solution. This is consistent with the
much smaller ionic radius of Mg compared with Ca.
(5) Tuite can accommodate the same or even higher
amounts of LILE and REE compared with apatite. Thus,
subsolidus apatite breakdown does not greatly affect LILE
or REE transport in subduction zones. A comparison of
element correlations indicates different mechanisms of
REE incorporation, namely Naþ þ REE3þ ¼ 2 Ca2þ for
tuite and Si4þ þ REE3þ ¼ Ca2þ þ P5þ for apatite.
(6) Only garnet can accommodate significant P in an
assemblage garnet þ omphacite þ SiO2 þ TiO2. In the
presence of apatite or tuite, P2O5 contents in garnet are in
the range of 0·2^0·6 wt % between 3 and 11 GPa and
increase to 0·8 wt % at 15 GPa in the absence of a
2057
JOURNAL OF PETROLOGY
VOLUME 50
detectable phosphate phase. This increase in P is associated
with the strongly increased solubility of Na and Ti (NaTimajorite component).
(7) The P storage capacity of clinopyroxene is limited to
250 ppm P and is, to a first approximation, independent
of its composition. Attempts to detect P in stishovite were
unsuccessful because of concentrations below the detection
limit of EMPA.
(8) Under subsolidus conditions apatite is likely to reach
its upper P stability limit in subducted MORB despite the
increasing P solubility in garnet. This is based on the bulk
P2O5 content of an average MORB and its modal composition along with the P storage capacity of eclogitic silicate
phases.
(9) As a result of the extreme preference of P for the
garnet structure, virtually the entire P budget of subducted
MORB will be locked up in garnet well into the lower
mantle, provided fO2 is high enough to prevent the stability of a metal phase.
AC K N O W L E D G E M E N T S
We are indebted to Hubert Schulze from BGI for his skilful
sample preparation and to Rainhard Kaindl (University
of Innsbruck) for assistance during Raman measurements.
Fabrice Brunet kindly provided samples of phosphorusdoped glass standards. Reviews by Fabrice Brunet, Cliff
Shaw and Ron Frost helped to correct various inaccuracies
and to improve the style of the manuscript. Their support
is gratefully acknowledged. This study was conducted
under the University of Innsbruck, Faculty of Geo and
Atmospheric Sciences’ research program ‘‘geodynamics^
geomaterials’’.
F U N DI NG
This work was supported by the Austrian Science
Foundation [grant number P14851-N04 to J.K.].
S U P P L E M E N TA RY DATA
Supplementary data are available at Journal of Petrology
online.
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