JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
PAGES 2067^2093
2012
doi:10.1093/petrology/egs044
An Experimental Study of Water in Nominally
Anhydrous Minerals in the Upper Mantle near
the Water-saturated Solidus
ISTVA¤N KOVA¤CS1,2*, DAVID H. GREEN1,3, ANJA ROSENTHAL1y,
JO«RG HERMANN1, HUGH ST. C. O’NEILL1, WILLIAM O. HIBBERSON1
AND BEATRIX UDVARDI2,4
1
RESEARCH SCHOOL OF EARTH SCIENCES, THE AUSTRALIAN NATIONAL UNIVERSITY, MILLS ROAD, BUILDING 61,
CANBERRA, ACT 0200, AUSTRALIA
2
DEPARTMENT OF DATA MANAGEMENT, GEOLOGICAL AND GEOPHYSICAL INSTITUTE OF HUNGARY, COLUMBUS U¤T
17^23, 1145, BUDAPEST, HUNGARY
3
SCHOOL OF EARTH SCIENCES AND CENTRE FOR ORE DEPOSIT STUDIES, UNIVERSITY OF TASMANIA, PTE. BAG 79,
HOBART, TASMANIA 7001, AUSTRALIA
4
LITHOSPHERE FLUID RESEARCH LAB, EO«TVO«S UNIVERSITY, PA¤ZMA¤NY PE¤TER SE¤TA¤NY 1/C, 1117, BUDAPEST,
HUNGARY
RECEIVED JULY 30, 2011; ACCEPTED JUNE 4, 2012
ADVANCE ACCESS PUBLICATION SEPTEMBER 2, 2012
The incorporation of water in olivine and pyroxenes interlayered
within fertile lherzolite compositions was explored experimentally
near the wet solidus of lherzolite at 2·5 and 4 GPa. The concentrations and activities of water were varied to establish the partitioning
of water between nominally anhydrous minerals (NAMs) and the
hydrous minerals pargasite and phlogopite. The water content in
NAMs was determined by Fourier-transform infrared (FTIR)
spectroscopy. The main absorption bands in NAMs from these experiments are very similar to those generally found in natural upper
mantle peridotites. Olivine, orthopyroxene and clinopyroxene contain
32^190, 290^320 and 910^980 ppm of water under the studied conditions. The partition coefficients between coexisting clinopyroxene
and orthopyroxene (Dcpx/opx) are 2·7 1·1 and 3·5 1·5 at 2·5
and 4 GPa respectively, whereas values for coexisting orthopyroxene
and olivine (Dopx/ol) are 6·7 2 and 4·7 1·1, at 2·5 and 4 GPa respectively.The storage capacity in NAMs in a model mantle composition close to the vapour-saturated solidus (water-rich vapour) is
190 ppm at both 2·5 and 4 GPa. Pargasite is the most important
phase accommodating significant amounts of water in the uppermost
mantle. Its breakdown with increasing pressure at 3 GPa at the
vapour-saturated solidus (which is at 10258C at 2·5 GPa) results
in a sharp drop in the water storage capacity of peridotite from
1000 ppm to 190 ppm H2O. At pressures 43 GPa, melting in
fertile lherzolite begins at the vapour-saturated solidus if the bulk
H2O concentration exceeds 190 ppm.
*Corresponding author: E-mail: [email protected]
yPresent address: Department of Earth Sciences, University of
Minnesota,108 Pillsbury Hall, Minneapolis, MN 55455, USA.
ß The Author 2012. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oup.com
KEY WORDS: upper mantle; partial melting; water; infrared spectroscopy; nominally anhydrous minerals; pargasite; phlogopite
I N T RO D U C T I O N
The identification of ‘water’ (i.e. H2O, OH, Hþ) as a
trace element at defect sites in nominally anhydrous
minerals (NAMs) has modified our understanding of
water storage in the mantle (Smyth et al., 1991; Bell &
Rossman, 1992). It is now well appreciated that water
in the NAMs of the Earth’s mantle has a major impact
on the mantle’s physical and chemical properties such as
VOLUME 53
NUMBER 10
OCTOBER 2012
et al., 1996; Rauch & Keppler, 2002; Stalder & Skogby,
2002; Berry et al., 2005, 2007; Bromiley et al., 2004; Smyth
et al., 2006; Grant et al., 2007b) are able to constrain the incorporation mechanisms of water in NAMs. Water incorporation in NAMs in peridotitic system has also been
studied by determining the partitioning between olivine,
orthopyroxene, clinopyroxene and a hydrous basaltic melt
in natural analogue starting compositions, using secondary ion-mass spectrometry (SIMS) (e.g. Koga et al., 2003;
Aubaud et al., 2004; Hauri et al., 2006; Tenner et al., 2009).
However, these experiments do not answer the allimportant question of how much water is held in the
NAMs of a given peridotitic mantle composition at saturation with a hydrous phase such as pargasitic amphibole
or phlogopite, or at the vapour-absent solidus. For example, although it is experimentally possible to measure
the water content of olivine at 2·5 GPa, 12008C in the presence of a water-rich vapour, under these conditions a lherzolite would be partially molten, and a vapour phase
would not be present except at extremely high bulk H2O
contents (45 wt %) (Green, 1973; Niida & Green, 1999;
Green et al., 2010, 2012; Fig. 1). There is no evidence that
such large concentrations of water occur in mantle rocks,
nor any compelling reasons to hypothesize that they might.
Maximum petrogenetic information is obtained from
melting experiments when a low degree of thermodynamic
4.0
3.5
3.0
w
at
fr er s
om a
t
G ur a
r e te
en d
(1 sol
97 id
3 ) us
melting temperature, viscosity, rheology, deformation pattern, elasticity and electrical conductivity.
For instance, the presence of water lowers the melting
temperature of mantle peridotite (e.g. Bowen, 1928;
Kushiro et al., 1968; Green, 1973; Milhollen et al., 1974;
Wyllie, 1978; Falloon & Danyushevsky, 2000). However,
there is disagreement as to how the solidus changes in the
presence of varying amounts of water for water-saturated
peridotite at changing pressure and temperature at uppermost mantle conditions owing to differences in experimental approaches and interpretations of the experimental
observations (Kushiro et al., 1968; Green, 1973; Milhollen
et al., 1974; Mysen & Boettcher, 1975; Green, 1976;
Wendlandt & Eggler, 1980; Mengel & Green, 1989;
Wallace & Green, 1991; Niida & Green, 1999; Grove et al.,
2006; Green et al., 2010, 2012). Water also lowers the viscosity of the mantle, facilitating its deformation and convection (Dixon et al., 2004). The deformation patterns of
mantle rocks change with changing water concentration
as different slip systems are activated at different concentrations (i.e. Karato et al., 1986; Karato & Wu, 1993;
Kaminski, 2002). It has been argued by using a rheological
modeling approach that water plays a substantial role in
the initiation of subduction and global plate tectonics
(Regenauer-Lieb et al., 2001; Regenauer-Lieb & Kohl,
2003; Li et al., 2008; Peslier et al., 2008).
Water concentration in NAMs seems to have a relatively
minor effect on both P and S seismic velocities (Karato,
1995; Jacobsen et al., 2004), but a larger effect on seismic
wave attenuation at seismic frequencies 51 Hz (Jackson
et al., 2002; Karato, 2006; Aizawa et al., 2008). The electrical conductivity of the mantle also depends on water
concentration (Karato, 1990, 2011; Wang et al., 2008), and
this property is seen by many as complementing the inferences derived from seismic methods in examining the
structure of the lower crust and upper mantle
(Gatzemeier & Moorkamp, 2005; Tommasi et al., 2006).
Although it is evident that water plays a fundamental
role in mantle processes, it has been difficult to assess the
mechanisms by which water is incorporated in mantle
minerals, and consequently how much water may be
accommodated. The direct measurement of water in
NAMs from mantle xenoliths (e.g. Bell & Rossman, 1992;
Grant et al., 2007a; Peslier, 2010) carries the ambiguity of
whether the ‘original’ mantle H2O contents are preserved
(Demouchy et al., 2006; Ingrin & Blanchard, 2006; Peslier
& Luhr, 2006; Peslier et al., 2008; Sundvall, 2010; Yang
et al., 2008). The fundamental problem is that the minerals
of xenoliths undergo subsolidus re-equilibration and are in
addition often metasomatized, so they do not carry direct
information on the water content of the typical mantle at
near-solidus conditions, which is a matter more appropriately addressed by experiment. Experiments in peridotitic
NAMs and chemically simple systems (e.g. Kohlstedt
Pressure (GPa)
JOURNAL OF PETROLOGY
2.5
solid+vapour
2 .0
solid+melt
1.5
1.0
0.5
900
1000
1100
1200
o
Temperature ( C)
Fig. 1. Experimental P^T conditions for model lherzolite compositions with mineral (i.e. NAMs) layers.
2068
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
variance is achieved and the chemical potentials of all
major components are constrained. For mantle melting
under hydrous conditions, the minimum variance condition is at the wet peridotite solidus where the four
anhydrous phases of normal peridotite (i.e. olivine þ
orthopyroxene þ clinopyroxene plus an aluminous phase,
plagioclase, spinel or garnet, depending on pressure), hydrous phases such as pargasite and/or phlogopite, melt
and vapour may all coexist. In this most interesting area
there were no experimental studies before the recent work
of Green et al. (2010, 2011) owing to the technical and analytical difficulties in measuring the water contents of
NAMs in equilibrated peridotitic bulk compositions under
controlled pressure, temperature and water conditions in
tiny experimental run products. Here we introduce a new
approach to the measurement of partitioning of water between melts, hydrous minerals and NAMs in equilibrium
(i.e. in terms of major elements and water) within peridotite under controlled pressure, temperature and bulk
water concentration, using Fourier transform infrared
(FTIR) spectroscopy.
Under the pressure^temperature conditions of the upper
mantle, the solubility of other components (i.e. Na2O,
K2O and SiO2) in an aqueous vapour phase reduces the
water fugacity in the vapour phase relative to pure water
at the same conditions (e.g. Bowen & Tuttle, 1949; Newton
& Manning, 2002; Dolejs & Manning, 2010). Furthermore,
in high-pressure peridotitic melting experiments, the partitioning of the alkali components Na2O and K2O into
such a vapour phase obviously depletes them in the other
phases of the system, particularly clinopyroxene, and, at
high water/rock ratios, may destabilize hydrous phases
such as pargasite or phlogopite (Green et al., 2010, 2011).
These changed phase compositions and stabilities alter
the melting behaviour of mantle lherzolite in ways that
depend on the amount of excess H2O, which explains the
apparent discrepancies between determinations of the ‘wet
solidus’ of mantle peridotite from different laboratories
(see Green, 1973; Grove et al., 2006; Green et al., 2010, 2011,
2012; Till et al., 2011).
In this study we use FTIR spectroscopy to determine
how the concentration of water in NAMs changes through
the water-saturated solidus at 2·5 and 4 GPa in fertile lherzolite compositions under controlled pressure, temperature
and bulk water concentrations. We also determine the
type of defects in which the water is stored. The method
uses layers of target phases placed in the experimental
charge to act as water sensors. The mineral grains formed
in the layers are large enough (minimum 30^50 mm) for
FTIR analysis, and sufficient randomly oriented grains
are available to make the statistical approach of Kova¤cs
et al. (2008) and Sambridge et al. (2008) feasible.
Green et al. (2010,2011,2012) clarified the roles of solute-rich
aqueous vapour, water-rich silicate melt, and pargasite and
phlogopite stability fields in a model mantle composition.
Green et al. (2010, 2011) found that the vapour-saturated solidus (water-rich vapour) of the lherzolite model mantle composition is 10108C at 2·5 GPa, 12108C at 4 GPa, and
13758C at 6 GPa. Inthis study, the technique of using‘melt-traps’ (layers of polycrystalline olivine and pyroxenes) facilitates the identification of solid phases quenched from
hydrous silicate melt and the distinction between hydrous
silicate melts and water-rich vapour.
E X P E R I M E N TA L A N D
A N A LY T I C A L M E T H O D S
General approach
Determining water contents of NAMs by FTIR spectroscopy has the advantage that the mechanism of water substitution may be identified from the spectra (e.g. Berry
et al., 2005), but quantification is dependent on calibrating
extinction coefficients for each substitution mechanism,
mineral composition and wavenumber (Paterson, 1982;
Bell et al., 1995, 2003; Libowitzky & Rossman, 1997; Sambridge et al., 2008; Kova¤cs et al., 2008, 2010). Alternative
techniques such as secondary ionization mass spectrometry (SIMS) determine the totals from all mechanisms of
water substitution but also unfortunately from any fluid inclusions present (i.e. Hauri et al., 2002; Koga et al., 2003;
Aubaud et al., 2007). Water in NAMs can be quantified by
analyzing 10^20 randomly oriented grains using unpolarized IR light (Kova¤cs et al., 2008; Sambridge et al., 2008).
It is not necessary to use large, oriented crystals, which
may be difficult to equilibrate with the matrix (i.e. Bai &
Kohlstedt, 1992; Zhao et al., 2004). Instead, crystals generated in the experiments can be utilized, although IR spectroscopy requires grain sizes of 20^100 mm, which are
generally not achieved in experimental subsolidus or
near-solidus runs. We have overcome this problem by
using sensor layers of olivine (Tables 1 and 2), which are
small enough to equilibrate completely with the neighbouring peridotite matrix during the experimental run
but large enough to be analysed by IR spectroscopy. An
additional advantage of using such monomineralic sensor
layers is that contamination of spectra by fluid inclusions
or water-rich phases can be avoided.
The experiments in which layers of olivine or discs of
single-crystal olivine were used as melt or vapour traps
(Table 2, Fig. 2) are from Green et al. (2010, 2011). The addition of olivine layers to the lherzolite composition does
not change the relative proportions of the key oxides such
as CaO, Al2O3, TiO2, Na2O, K2O and H2O but may
slightly alter Mg# [100Mg/(Mg þ Fe)] if the Mg# of
the added olivine differs from that of the lherzolite (Green
et al., 2010). Additional experiments were performed specifically for this study in which polycrystalline orthopyroxene
and clinopyroxene were added as layers either with or
2069
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Table 1: Starting composition (wt %) of lherzolite mixes and pyroxenes
HZ1
SiO2
46·2
HZ2
46·8
HZ1 þ 5 wt % ‘dry’ phl
Al-poor
46·0
Al-rich
enstatite
chrome-diopside
enstatite
diopside
DTP 11
D25011
E25542
D25542
56·2
53·3
53·2
49·7
TiO2
0·18
0·19
0·17
0·01
0·09
0·19
0·31
Al2O3
4·08
4·28
4·50
2·85
2·59
6·30
6·80
FeO*
7·59
4·14
7·18
3·81
1·57
5·80
2·38
Fe2O3
n.a.
n.a.
n.a.
0·43
0·04
0·74
1·37
MnO
MgO
0·10
38·0
0·10
39·8
0·10
0·09
37·5
34·2
0·05
16·6
0·03
17·8
2·11
20·4
CaO
3·23
3·39
3·05
1·15
Na2O
0·33
0·35
0·31
0·01
1·39
0·06
K2O
0·03
0·03
0·62
0·01
0·00
0·02
0·05
P2O5
n.a.
n.a.
n.a.
0·04
0·04
n.a.
n.a.
Cr2O3
0·40
0·42
0·86
1·41
0·67
0·89
NiO
0·28
0·29
0·27
n.a.
n.a.
n.a.
n.a.
H2O (ppm)3
n.a.
n.a.
n.a.
22·4
0·11
31·21
0·36
49
83
110
Total
100·4
99·8
100·1
99·6
99·6
100·4
1175
100·0
Mg#
89·9
94·5
90·3
93·6
94·9
88·6
89·8
*For HZ1 and HZ2 mix iron expressed as total iron in FeO.
1
A. J. Easton’s analysis (1963).
2
From Green (1964).
3
Water content is determined by FTIR using the calibration factors of Bell et al. (1995). The analytical uncertainty is
c. 30%.
n.a., not analysed.
without olivine (Table 2, Fig. 2; Green et al., 2010), to quantify the amounts of water stored in these phases at the
multiply saturated solidus. Pyroxenes have significant
SiO2, CaO, Al2O3, TiO2 and Na2O contents and these
components exchanged readily with the lherzolite layer.
Thus these experiments were not used by Green et al.
(2010, 2011) to determine the phase equilibria of the lherzolite (HZ composition).
Experimental procedures
Details of experimental methods have been given by Green
et al. (2010). Here we focus on the experimental and analytical procedures used to quantify water incorporation and
substitution mechanism in peridotitic NAMs. Two lherzolitic compositions were investigated, one matching the
upper mantle composition of Hart & Zindler (1986), with
Mg# ¼ 89·9 (HZ1, Table 1), and the other being more
magnesian (Mg# ¼ 94·5; HZ2, Table 1), with approximately half the FeO content of HZ1. Each composition
was prepared in duplicate as described by Green et al.
(2010), both as an anhydrous mix, using fired oxides
including MgO, and as a hydrous mix by substituting
Mg(OH)2 for the MgO, giving water contents of 0 and
14·5 wt % H2O respectively. These end-member compositions were combined in appropriate proportions to give
starting mixes with water contents of 2·9, 1·45 and 0·145 wt
%. The roles of K2O and phlogopite were investigated by
adding 5 wt % of a ‘dry’ phlogopite component to the
HZ1 composition; that is, HZ1 þ5 wt % ‘dry’ phlogopite
(Table 1). In addition to the experiments with controlled
water contents, two ‘dry’ experiments were conducted
below the solidus, one at 2·5 GPa without, and one at
4·0 GPa with, the 5 wt % ‘dry’ phlogopite component.
The mixes were loaded in silver, gold or, in one experiment, gold^palladium (Au25Pd75) capsules, with polycrystalline layers of crushed natural crystals of olivine,
orthopyroxene or clinopyroxene at the top and bottom of
the capsules (Table 2). Five kinds of crystals were used:
Al-rich orthopyroxene (E2554) and clinopyroxene
(D2554) from the primary assemblage of the Lizard
Peridotite, Cornwall (Green, 1964); Al-poor chromediopside clinopyroxene (D2501) from a garnet lherzolite
from Almklovdalen, Western Norway; Al-poor orthopyroxene (enstatite, DTP1) from a garnet harzburgite xenolith
2070
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
Table 2: Experimental conditions and results on HZ1, HZ2 and 95% HZ1 þ 5% ‘dry phlogopite’compositions
Run no. Mount no. T (8C) Time
wt % H2O Capsule Top L
Bottom L
(days)
Phase assemblage
‘Water’ in ‘Water’ in
(with OL þ OPX þ CPX þ GA)
OL (ppm) OPX (ppm) CPX (ppm)
PHL
PAR
Melt
‘Water’ in
Vap
HZ1 peridotite at 2·5 GPa
D968
P30
1000
7
DRY
Au
OL(SC)
Lo-Al Opx —
PAR
—
—
32
D897
O85
1000
2·7
1·45
Ag
OL(SC)
OL(SC)
PAR
—
Vap
67
D937
P4
1000
7
1·45
Au
Hi-Al Opx Hi-Al Cpx —
—
—
Vap
472
1431
D944
P7
1000
7
1·45
Au
Lo-Al Opx Lo-Al Cpx —
—
—
Vap
365
978
C2936
O79
1025
3
1·45
Ag
OL(SC)
OL(SC)
—
—
Melt
Vap
71
C2930
O77
1050
3
1·45
Ag
OL(SC)
OL(SC)
—
—
Melt
—
68
35
—
224
HZ2 peridotite at 2·5 GPa
C2888
O60
1000
3
2·9
Ag
OL-Disc
—
PAR
—
Vap
C2886
O52
1025
3
1·45
Ag
OL(SC)
OL(SC)
—
PAR
—
Vap
31
C2887
O53
1050
3
1·45
Ag
OL(SC)
OL(SC)
—
—
Melt
—
52
95% HZ1 þ 5% ‘anhydrous phlogopite’ at 2·5 GPa
D949
P10
1000
3
0·145
Au
OL(SC)
Lo-Al Opx PHL
PAR
—
—
44
291
D948
P9
1000
3
1·45
Au
OL(SC)
Lo-Al Opx PHL
PAR
—
Vap
44
298
HZ1 peridotite at 4·0 GPa
C2987
O92
1100
4
0·145
Au
OL(SC)
OL(SC)
—
—
—
Vap
110
C2942
O80
1100
3
1·45
Au
OL(SC)
OL(SC)
—
—
—
Vap
188
C3005
P3
1150
7
1·45
Au
Hi-Al Opx Hi-Al Cpx —
—
—
Vap
377
1590
C3010
P6
1150
7
1·45
Au
Lo-Al Opx Lo-Al Cpx —
—
—
Vap
260
910
C2950
O81
1200
3
1·45
Au
OL(SC)
OL(SC)
—
—
—
Vap
C2899
O98/99
1225
1
1·45
AuPd
OL(SC)
OL(SC)
—
—
Melt
—
30
Ol-Disc
—
—
—
Vap
80
OL(SC)
—
—
—
Vap
62
115
HZ2 peridotite at 4·0 GPa
C2889
O61
1100
3
2·9
Ag
C2928
O76
1150
1
1·45
Ag
OL(SC)
95% HZ1 þ 5% ‘anhydrous phlogopite’ at 4·0 GPa
C3087
P32
1175
1
DRY
Au
OL(SC)
Lo-Al Opx PHL
—
—
—
57
264
C3029
P12
1100
4
0·145
Au
OL(SC)
Lo-Al Opx PHL
—
—
—
65
299
C3024
P11
1100
3
1·45
Au
OL(SC)
Lo-Al Opx PHL
—
—
Vap
71
317
C3014
P8
1150
3
1·45
Au
OL(SC)
Lo-Al Opx —
—
Melt
—
62
311
OL(SC), olivine (San Carlos); OPX, orthopyroxene; CPX, clinopyroxene; GA, garnet; PAR, pargasite; PHL, phlogopite;
Vap, vapour; L, Layer; Lo-Al OPX, low-alumina orthopyroxene; Hi-Al OPX, high-alumina orthopyroxene; AuPd, AuPd
double capsule. The maximum uncertainty in the reported concentrations is 30%.
in the Dutoitspan kimberlite, South Africa; San Carlos
olivine (Fig. 3). Major element analyses of these pyroxenes
are given in Table 1.
We deployed four configurations: configuration 1, San
Carlos olivine (SC) layers at the top and bottom, using
HZ1 or HZ2 bulk compositions; configuration 2, Al-rich
clinopyroxene (D2554) at the top and Al-rich orthopyroxene (E2554) at the bottom, using fertile HZ1 lherzolite
bulk composition; configuration 3, Al-poor clinopyroxene
(D2501) at the top and Al-poor orthopyroxene (DTP1) at
the bottom, using fertile HZ1 lherzolite bulk composition;
configuration 4, San Carlos olivine (SC) at the top and
Al-poor orthopyroxene (DTP1) at the bottom, using fertile
HZ1 lherzolite or HZ1 þ5 wt % ‘dry’ phlogopite bulk compositions (Fig. 2, Table 2). Two experiments (O60, O61,
Table 2) were carried out at the start of the campaign
with single-crystal discs of San Carlos olivine within the
lherzolite-filled capsule, but because the recovered crystals
contained many fluid inclusions along healed fractures,
which dominated the FTIR spectra, this approach was
not continued, and crushed minerals were used instead.
The olivine-disc experiments nevertheless served to
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JOURNAL OF PETROLOGY
NUMBER 10
Lo-Al Opx (DTP1) layer
Ol(SC) layer
3.5 mm
San Carlos olivine layer
VOLUME 53
HZ1 fertile lherzolite layer
+ 5 wt % ‘dry’ phl
OCTOBER 2012
OL(SC)
HZ1 + 5 wt % ‘dry’ phl
Lo-Al opx (DTP1)
Lo-Al orthopyroxene (DTP1) layer
ol
ol
ol
opx
500 µm
500 µm
opx
500 µm
Fig. 2. Photomicrograph of the P12 experimental charge with San Carlos olivine (SC) and Al-poor orthopyroxene (DTP1) layers in the
K-enriched HZ1 fertile lherzolite mix. The images are taken in plane-polarized light.
40
Al-rich orthopyroxene (E2554)
Al-poor orthopyroxene (DTP1)
35
25
Al-rich clinopyroxene (D2554)
20
Al-poor clinopyroxene (D2501)
15
Absorbance/cm
30
10
San Carlos olivine
5
3800 3700 3600 3500 3400 3300 3200 3100 3000
Wavenumber (cm-1)
0
3700 3600 3500 3400 3300 3200 3100 3000
Wavenumber (cm-1)
Fig. 3. Average unpolarized IR spectra for the original pyroxenes and San Carlos olivine, normalized to 1cm thickness, used as starting minerals interlayered within model lherzolite compositions. Dashed lines represent the spectra of original ‘sensor’ minerals. Continuous lines indicate the spectra of the respective ‘sensor’ minerals in some experiments at 4 GPa and 11008C. (See Figs 4^6 and text for details.) Spectra are
stacked for clarity.
2072
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
demonstrate the presence of vapour rather than silicate
melt and the ability of the vapour phase to react with
olivine along microfractures within the olivine discs
(Experiment O89, Table 5) (Green et al., 2010).
Experiments were conducted in end-loaded pistoncylinder apparatuses at 2·5 and 4 GPa and temperatures
of 1000^10508C and 1100^12258C respectively (Table 2;
Green et al., 2010). The Au, Ag or Au25Pd75 capsules were
placed within NaCl^Pyrex sleeves, a cylindrical graphite
furnace, and internal spacers of crushable MgO. Oxygen
fugacity is not buffered and may vary in experiments, as
it is a function of H2O activity, furnace assembly components, any Fe loss to the capsules, and the starting material.
Oxygen fugacity, however, is inferred to lie between the
fayalite^magnetite^quartz (FMQ) and iron^wu«stite (IW)
buffers based on the results of Niida & Green (1999) using
similar furnace assemblies, and calculation of fO2 from
spinel stoichiometry. In higher temperature experiments
where it was necessary to use AuPd double capsules, the
mineralogy indicates higher oxidation state (see below).
Pressure was calculated from the direct conversion of load
to pressure (no friction correction) and is accurate to
0·1GPa. The experiments ran from 1 to 7 days, with
longer durations used in particular for lower temperature
and nominally ‘dry’ runs to promote the attainment of
equilibrium. Temperature was controlled to an estimated
accuracy of 108C and precision of 28C, using a
Eurotherm 904 controller affixed to type-B thermocouple
(Pt94Rh6/Pt70Rh30). The recovered samples were mounted
in epoxy and polished after exposure of a representative
section.
Analytical methods
The phase compositions, phase relations and grain sizes
were determined by energy-dispersive spectrometry
(EDS) using a JEOL 6400 scanning electron microscope
(SEM), and additional analyses and imaging were performed using a Hitachi 4300 field emission SEM
(FESEM), both operating at 15 kV and a beam current
of 1 nA. All facilities are housed in the Electron
Microscopy Unit (EMU) of the ANU. Mineral standards produced by Astimex Scientific Limited were used
to standardize mineral and glass analyses. Detection
limits are 0·10 wt % for K2O, TiO2 and MnO,
0·15 wt % for Cr2O3, and 0·15 wt % for Na2O.
Analyses by EDS^SEM have advantages in comparison
with wavelength-dispersive spectrometry (WDS)^microprobe in allowing analysis of fine-grained experimental
run products at low beam current (minimizing element
volatilization, particularly Na) and simultaneous rather
than sequential analyses for the selected elements. The
accuracy of the EDS^SEM methods used in this and
similar studies from ANU, relative to the WDS^microprobe methods has been demonstrated (Spandler et al.,
2010) with multiple analyses of garnet, clinopyroxene
and plagioclase.
After SEM analysis, doubly polished thin-sections were
made for FTIR analysis, with thicknesses from 37 to
124 mm (Tables 3 and 4). The thickness of the doubly polished section was measured with a Mitotuyo analogue micrometer, which is nominally accurate to within 2 mm. A
Bruker IFS-28 IR spectrometer mounted with an A590
Bruker IR microscope, supplied with a nitrogen-cooled
MCT detector, and a KBr beam splitter was used for IR
analysis (see Berry et al., 2005; Kova¤cs et al. 2008, 2010,
for further details). Spectra were recorded in the range
600^5000 cm1. The spectra have a resolution of 2 cm1.
Analyses were made with a circular aperture of
30^100 mm diameter (depending on the target grain’s size)
while the microscope stage was continuously flushed with
nitrogen. Spectra were processed using the OPUSÕ software (Bruker Inc.). For background subtraction the ‘interactive concave rubberband correction’ (ICRC) tool within
the OPUSÕ software was used where there was a relatively
smooth background, but it was drawn manually where it
was irregular owing to water vapour, fluid inclusions, etc.
The alternative background correction routines provide
similar integrated areas within 5%. The integrated intensities of the main absorption bands were obtained with
the Integration tool of the OPUSÕ software using the integration limits given in Electronic Appendix Table EA1
(available for downloading at http://www.petrology
.oxfordjournals.org). The total absorbance was calculated
from the average unpolarized spectra. The precision in
the total integrated transmittance, and also, absorbance
[i.e. as theoretically shown by Sambridge et al. (2008) and
tested by Kova¤cs et al. (2008)], normalized to unit thickness, which is simply proportional to water content, is subject to (1) uncertainty in the measured integrated
absorbance from each spectrum, and in the thickness of
the sample, and (2) uncertainty in the estimation of the
total absorbance, which depends on the number of grains
analysed for the average unpolarized absorbance (e.g.
Sambridge et al., 2008). Here we have at least nine grains,
suggesting that this error should be510 %. In addition
(see below for further details), the calibration factors also
introduce an error, thought to be less than 15% (Bell
et al., 1995; Kova¤cs et al., 2010). If all of these factors are considered, a maximum error of 30% in each datum appears
to be realistic, but should typically be lower than this.
There are major substitution mechanisms for water in
olivine, associated with (1) Si vacancies, (2) Mg vacancies,
(3) octahedral Ti, and (4) trivalent cations compensated
by H bound to oxygens at the tips of the vacancies,
which hereafter we label as [Si] (3450^3630 cm1), [Mg]
(3100^3300 cm1), [Ti] (3525 and 3572 cm1) and [triv]
(3300^3400 cm1), respectively (Berry et al., 2005, 2007;
Walker et al., 2007; Kova¤cs et al., 2010; Balan et al., 2011).
2073
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Table 3: Infrared characteristics of olivine in chemically complex experiments
t (mm)
[Si þ Ti]
n.o.a.
[triv]
H2O (ppm)
K*
Bands (cm1):
3613 þ
3572 [Ti]
3525 [Ti]
3480 [Si]
3450 [Si]
3598 [Si]
B
3354 þ
3329 [triv]
O61
75(5)
15
170
40
71
46
4
8
59
229
80
O76
83(2)
19
143
7
86
46
3
1
76
218
62
41
O80
77(2)
12
422
129
192
73
19
9
122
544
188
102
O92
124(5)
17
257
47
135
66
7
2
87
344
110
65
O81
71(5)
19
211
65
99
33
9
5
185
396
115
75
O98
77(4)
12
23
0
20
0
0
3
139
161
30
30
Bands (cm1):
3611 [Si]
3597 þ
3567y [Ti]
3546 [Si]
3526 [Ti]
3480 [Si]
3448 [Si]
3588 [Si]
43
3354 þ
3329 [triv]
O60
70(5)
25
130
11
5
41
7
60
0
6
1
131
35
25
O85
83(3)
9
146
42
13
53
5
23
6
5
70
217
67
41
O52
73(5)
12
91
9
0
45
0
33
3
1
52
143
31
27
O79
71(2)
11
155
46
10
58
4
26
7
4
86
241
71
45
O53
63(5)
14
129
27
7
53
2
33
4
2
68
196
52
37
O77
80(5)
11
140
40
14
50
6
19
5
7
86
226
68
42
3613 [Si]
3596 þ
3568y [Ti]
3545 [Si]
3525 [Ti]
3480 [Si]
3450 [Si]
3354 þ
Bands (cm1):
3587 [Si]
3329 [triv]
P8
74(9)
15
110
41
7
45
10
4
0
3
102
212
62
40
P11
65(2)
15
137
39
11
58
6
8
7
10
104
241
71
45
P12
70(3)
15
136
39
8
61
3
11
10
4
86
222
65
42
P9
60(2)
15
81
20
10
33
3
7
4
5
74
155
44
29
P10
67(4)
15
81
21
4
34
4
9
5
5
80
161
44
30
P32
81(5)
13
83
24
4
36
5
6
3
5
148
230
57
43
P30
72(5)
12
76
17
2
41
1
11
1
3
50
127
32
24
Absorbances are in integrated total absorbance normalized to 1 cm thickness. t, thickness, with standard deviation given
in parentheses. n.o.a., number of analyses; [Si], silica-vacancy substitution; [Ti], Ti-clinohumite substitution; [triv], trivalent cation-related substitution (see text for details).
yThese bands more likely represent [Ti] peaks or a combination of [Ti] and [Si] bands. K, Kovács et al. (2010); B, Bell
et al. (2003). k ¼ 0·188.
*Calibration factors of 0·572, 0·182, and 0·178 are applied to [Si], [Ti] and [trivalent] bands respectively (see Kovács et al.,
2010).
Quantifying water contents from IR absorption spectra requires the use of calibration factors, and Kova¤cs et al.
(2010) pointed out that these four common substitution
mechanisms for water in olivine each requires its own
calibration factor: k[Si] ¼ 0·57 0·04, k[Ti] ¼ 0·18 0·07,
k[triv] ¼ 0·18 0·05, and k[Mg] ¼ 0·03 0·03. Recently,
Balan et al. (2011) provided additional theoretical support
for the different substitution mechanisms requiring different calibration factors.
The challenge is to resolve accurately the contribution of
the [Ti] and [Si] bands between 3500 and 3600 cm1.
The [Ti] mechanism can be identified by two major
bands at 3572 and 3525 cm1 and where these were visible
we used the calibration factor for this mechanism, as the
amounts of water associated with [Si] are probably low at
the relative high silica activity of the experiments. For the
sake of comparison, the water concentrations for olivine
calculated by the Bell et al. (2003) calibration are also
2074
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
Table 4: Infrared characteristics and parameters for orthopyroxene and clinopyroxene
Sample
Thickness (mm)
n.o.a.
Atot/cm
H2O (ppm)
Bell et al. (1995)
Orthopyroxene
Configurations 2 and 3
P3
60(2)
10
5590
377
P4
67(4)
10
7010
472
P6
53(3)
10
3856
260
P7
37(4)
10
5418
365
P8
74(9)
10
4620
311
P9
65(2)
10
4428
298
P10
70(3)
10
4322
291
P11
60(2)
10
4702
317
P12
67(4)
10
4438
299
P30
81(5)
10
3329
224
P32
72(5)
12
3923
264
E2554
144(13)
10
544
110
DTP1
161(21)
20
242
49
Configuration 4
Clinopyroxene
Configurations 2 and 3
P3
60(2)
15
11275
1590
P4
67(4)
15
10151
1431
P6
53(3)
15
6453
910
P7
37(4)
15
6938
978
D2554
165(2)
10
2777
1175
D2501
191(20)
15
196
83
Values in parentheses are standard deviations. The absorbances are given in integrated total absorbance normalized
to 1 cm thickness. n.o.a., number of unpolarized FTIR
analyses.
reported in Table 3, and these usually give lower values
than the Kova¤cs et al. (2010) calibration, especially when
the contribution of the [Si] bands is significant.
For the pyroxenes the calibration factors of Bell et al.
(1995) were applied. The mechanism by which water substitutes in both orthopyroxene and clinopyroxene is less variable than in olivine, producing spectra with roughly the
same absorption characteristics in all experiments. These
spectra are similar to those in the samples used for calibration by Bell et al. (1995), showing mainly high wavenumber
bands at 43400 cm1; hence the calibration factors of Bell
et al. (1995) were taken to be applicable.
R E S U LT S
Achievement of equilibrium
Analyzing the NAM sensor crystals before and after
experiments both for major elements and water allows us
to check whether they approached equilibrium in both respects. The original San Carlos olivine has Mg# 90·5
according to Galer & O’Nions (1989) although samples
display a small variation in colour. Olivine grains are
70^120 mm in diameter in the layers and chemically homogeneous in each run, with the Mg# varying from 90 to
91·3 for olivine coexisting with the HZ1 composition
(Table 5). However, the interlayered olivine becomes
more magnesian (Mg# ¼ 91·5^93·2) in the HZ2 lherzolite
mix (Table 5). The absorption characteristics of O^H vibration bands of San Carlos olivine also changed as [triv]
(3329 and 3354 cm1) and [Si] (3598 and 3612 cm1)
bands appeared in addition to or replacing the original
[Ti] bands (3525 and 3572 cm1) in all studied experimental configurations (Fig. 3). Inhomogeneity of water concentration was not detected in the sensor minerals before and
after the experiments (no diffusion profiles could be
observed).
Pyroxenes of configurations 2, 3 and 4 changed their
chemical composition by exsolving an aluminous phase
(garnet), which lowers their Al2O3 content with respect to
the starting pyroxenes by adjusting to the new bulk compositions (Table 5; Green et al., 2010; see below for details).
In the IR spectra of pyroxenes, both the position and contribution of the original bands changed with respect to
the starting pyroxenes (Fig. 3). The original Al-rich orthopyroxene (E2554) has three major absorption bands at
3565, 3525 and 3420 cm1 contributing 110 ppm of water
(Fig. 3, Table 1). The original Al-poor orthopyroxene
(DTP1), in contrast, has its major absorption band at
3600 cm1 with a smaller one at 3420 cm1, and only
49 ppm water (Fig. 3, Table 1). The spectra of newly
formed orthopyroxene in both configuration 2 and 3 experiments all have major, broad bands at 3600, 3530 and
3420 cm1 clearly differing from the original orthopyroxene (Fig. 3). The original Al-rich clinopyroxene (D2554)
has 1180 ppm of water, with three major absorption bands
at 3565, 3525 and 3420 cm1 and a broad band at
3675 cm1 (Fig. 3, Table 1). The broad band at 3675 cm1 is
typical of hydrous minerals such as amphibole. The original Al-poor clinopyroxene (D2501), in contrast, has a
much lower concentration of water (83 ppm) and two
major bands at 3540 and 3450 cm1 (Fig. 3, Table 1). The
re-equilibrated clinopyroxene in all configuration 2 and 3
experiments has two broad bands at 3640 and 3450 cm1
and a smaller one at 3360 cm1 (Fig. 3) regardless of
whether the starting clinopyroxene was Al-rich or
Al-poor, indicating that equilibrium was approached in
the experiments for the water substitutions. There is no evidence that significant water defects are inherited from the
original ‘sensor’ crystals (Fig. 3). Instead, new substitution
mechanisms formed (i.e. additional bands appeared) and
the original bands decreased in intensity, broadened, narrowed or vanished during the experiments (Fig. 3).
2075
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
Table 5: Major element composition of minerals in the sensor layers and the lherzolite mix
Phase
n
SiO2 1s
TiO2 1s
Al2O3 1s
Cr2O3 1s
NiO 1s
FeO 1s
MgO 1s
CaO 1s
Na2O 1s
K2O 1s
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
Total
Mg#
bdl
bdl
bdl
0·58 0·09
5·73 0·18
53·37 0·55
bdl
bdl
bdl
102·58
94·3
HZ2 Peridotite at 2.5 GPa
Run O-60 (10008C, 2·9% H2O)*
OL Lhz
5
42·90 0·40
Run O-89 (10008C, 7·25% H2O)*
OL Lhz
5
42·35 0·51
bdl
bdl
bdl
0·49 0·16
6·13 0·46
52·52 0·34
bdl
bdl
bdl
101·48
93·9
OL disc (central)
6
41·82 0·42
bdl
bdl
bdl
0·31 0·15
8·40 0·22
50·59 0·48
bdl
bdl
bdl
101·11
91·5
OL disc (edge)
5
42·66 0·36
bdl
bdl
bdl
0·58 0·07
6·06 0·17
52·13 0·63
bdl
bdl
bdl
101·42
93·9
Run O-52 (10258C, 1·45% H2O)
OL Lhz
3
41·51 0·43
bdl
bdl
bdl
0·49 0·09
6·43 0·38
50·86 1·51
bdl
bdl
99·39
93·4
OL-L
5
42·75 0·33
bdl
bdl
bdl
0·37 0·15
7·51 0·48
52·45 0·74
bdl
0·12 0·05
bdl
bdl
103·08
92·6
Run O-53 (10508C, 1·45% H2O)
OL Lhz
2
42·74 0·17
bdl
bdl
bdl
0·51 0·08
6·45 0·69
51·82 0·44
bdl
bdl
bdl
101·52
93·5
OL-L
2
42·37 0·49
bdl
bdl
bdl
0·46 0·09
8·40 0·66
50·37 0·11
bdl
bdl
bdl
101·59
91·5
bdl
HZ2 Peridotite at 4·0 GPa
Run O-61 (11008C, 2·9% H2O)*
OL Lhz
2
42·18 0·04
bdl
bdl
bdl
0·52 0·04
5·81 0·04
52·59 0·08
OL disc-edge
1
41·91
bdl
bdl
bdl
0·28
6·72
51·73
bdl
bdl
101·09
94·2
0·20
bdl
bdl
100·84
93·2
0·13 0·06
bdl
bdl
101·08
93·3
bdl
bdl
99·49
92·4
Run O-76 (11508C, 1·45% H2O)
OL Lhz
2
41·98 0·35
bdl
bdl
bdl
0·39 0·15
6·63 0·18
51·96 0·31
OL-L
1
41·40
bdl
bdl
bdl
0·23
7·39
50·47
bdl
HZ1 Peridotite at 2·5 GPa
Run P-30 (10008C, dry)
OL Lhz
7
41·65 0·94
bdl
bdl
bdl
0·32 0·24
9·53 0·38
50·29 0·78
bdl
bdl
bdl
101·78
90·4
OL-L
3
41·67 0·27
bdl
bdl
0·18 0·08
0·16 0·21
9·00 0·56
50·19 0·98
bdl
bdl
bdl
101·20
90·9
OPX Lhz
8
56·78 0·84
bdl
2·08 0·52
0·21 0·13
bdl
6·30 0·40
34·35 0·58
0·90 0·23
bdl
bdl
100·62
90·7
OPX-L
6
57·19 0·66
bdl
2·34 0·22
0·82 0·11
bdl
4·57 0·26
35·61 0·21
0·62 0·07
bdl
bdl
101·14
93·3
Run O-85 (10008C, 1·45% H2O)
OL Lhz
3
41·79 0·70
bdl
bdl
bdl
0·50 0·32
9·86 0·24
50·00 0·89
bdl
bdl
bdl
102·15
90·0
OL-L
1
41·77
bdl
bdl
0·18
0·61
9·86
49·75
bdl
bdl
bdl
102·17
90·0
91·1
Run P-4 (10008C, 1·45% H2O)
OPX Lhz
4
55·70 0·92
0·16 0·04
2·50 0·21
0·41 0·06
0·19 0·18
5·78 0·11
33·32 0·64
0·72 0·06
0·21 0·27
bdl
98·97
OPX-L
6
55·82 1·82
bdl
3·39 0·90
0·71 0·18
0·18 0·07
6·12 0·33
33·03 1·39
0·88 0·16
bdl
bdl
100·12
90·6
CPX Lhz
8
52·88 1·08
0·44 0·09
3·25 0·45
1·01 0·19
0·13 0·22
2·82 0·15
17·14 0·44
20·81 0·57
0·94 0·22
bdl
99·42
91·6
CPX-L
5
52·06 1·24
0·39 0·06
4·38 1·24
0·97 0·21
bdl
2·54 0·21
16·71 0·88
21·51 0·97
0·68 0·19
bdl
99·23
92·1
Run P-7 (10008C, 1·45% H2O)
OPX Lhz
5
57·21 0·40
bdl
2·42 0·42
0·56 0·13
0·18 0·18
5·52 0·22
34·38 0·20
0·76 0·07
bdl
bdl
101·02
91·7
OPX-L
4
56·75 0·42
bdl
2·53 0·19
0·83 0·06
bdl
4·71 0·14
34·79 0·47
0·73 0·06
bdl
bdl
100·34
92·9
CPX Lhz
7
53·60 0·43
0·25 0·11
3·43 0·39
1·08 0·24
bdl
2·75 0·28
17·38 0·40
20·81 0·23
1·12 0·33
bdl
100·42
91·9
CPX-L
6
54·38 0·48
0·21 0·06
2·83 0·56
1·47 0·26
bdl
1·77 0·24
16·60 0·56
21·46 0·22
1·25 0·44
bdl
99·95
94·4
bdl
bdl
bdl
0·51 0·11
9·46 0·06
49·39 0·14
0·12 0·03
bdl
bdl
100·81
90·3
Run O-79 (10258C, 1·45% H2O)
OL Lhz
4
41·34 0·40
Run O-77 (10508C, 1·45% H2O)
OL Lhz
5
41·37 0·46
bdl
bdl
bdl
0·39 0·14
9·67 0·12
49·22 0·52
bdl
bdl
bdl
100·65
90·1
OL-L
2
40·90 0·18
bdl
bdl
bdl
0·37 0·22
9·74 0·19
48·90 0·12
bdl
bdl
bdl
99·90
90·0
HZ1 Peridotite at 4·0 GPa
Run O-92 (11008C, 0·145% H2O)
OL Lhz
4
41·50 0·39
bdl
bdl
bdl
0·37 0·09
9·31 0·14
49·78 0·37
bdl
bdl
bdl
100·95
90·5
OL-L
2
41·42 0·07
bdl
bdl
bdl
0·30 0·31
9·23 0·07
50·07 0·03
bdl
bdl
bdl
101·02
90·6
Run O-80 (11008C, 1·45% H2O)
OL Lhz
5
41·76 0·05
bdl
bdl
bdl
0·36 0·18
9·37 0·22
50·17 0·37
bdl
bdl
bdl
101·65
90·5
OL-L
1
42·10
bdl
bdl
bdl
0·40
8·92
50·04
bdl
bdl
bdl
101·46
90·9
bdl
1·40 0·20
0·28 0·05
bdl
5·50 0·06
34·42 0·41
bdl
bdl
100·10
91·8
Run P-3 (11508C, 1·45% H2O)
OPX Lhz
5
57·44 0·46
1·06 0·09
(continued)
2076
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
Table 5: Continued
Phase
n
SiO2 1s
TiO2 1s
Al2O3 1s
Cr2O3 1s
NiO 1s
FeO 1s
MgO 1s
CaO 1s
Na2O 1s
K2O 1s
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
[wt%]
Total
Mg#
OPX-L
4
57·74 0·15
bdl
1·40 0·15
0·38 0·11
bdl
5·40 0·19
34·72 0·27
0·97 0·13
bdl
bdl
100·60
92·0
CPX Lhz
5
55·00 0·17
0·12 0·10
1·85 0·24
0·56 0·14
bdl
3·03 0·29
19·24 0·22
19·87 0·83
0·70 0·16
bdl
100·36
91·9
CPX-L
7
54·76 0·46
0·16 0·13
1·98 0·32
0·57 0·11
bdl
2·61 0·17
18·60 0·60
21·22 0·59
0·61 0·07
bdl
100·50
92·7
Run P-6 (11508C, 1·45% H2O)
OPX Lhz
6
57·65 0·39
bdl
1·57 0·18
0·43 0·07
0·15 0·08
5·05 0·14
34·85 0·26
1·04 0·09
bdl
bdl
100·73
92·5
OPX-L
6
57·26 0·52
0·11 0·07
1·57 0·15
0·48 0·15
bdl
4·54 0·18
34·69 0·29
0·89 0·06
bdl
bdl
99·52
93·2
CPX Lhz
CPX-L
12
55·02 0·48
0·11 0·07
2·42 0·15
0·90 0·21
bdl
2·73 0·35
18·87 0·40
19·16 0·42
1·16 0·09
bdl
100·37
92·5
6
53·61 0·96
0·17 0·08
2·44 0·11
1·39 0·16
bdl
1·93 0·21
17·58 0·39
19·54 0·39
1·32 0·08
bdl
97·97
94·2
bdl
bdl
bdl
0·19 0·22
8·48 0·56
49·82 0·93
bdl
bdl
bdl
99·62
91·3
Run O-81 (12008C, 1·45% H2O)
Ol Lhz
5
41·12 0·50
Run O-98/99 (12258C, 1·45% H2O)
OL Lhz
4
41·79 0·33
bdl
bdl
bdl
0·38 0·15
9·06 0·35
50·21 0·61
0·11 0·05
bdl
bdl
101·55
90·8
OL-L
5
41·92 0·26
bdl
bdl
bdl
0·41 0·20
8·96 0·29
50·47 0·32
0·14 0·08
bdl
bdl
101·90
90·9
95% HZ 1þ 5% ’anhydrous phlogopite’ at 2·5 GPa
Run P-10 (10008C, 0·145% H2O)
Ol Lhz
4
41·61 0·60
bdl
0·28 0·42
bdl
0·32 0·01
9·18 0·43
50·23 0·68
bdl
bdl
101·61
90·7
OPX Lhz
3
57·68 0·60
bdl
2·04 0·32
0·31 0·20
0·19 0·10
5·68 0·36
34·86 0·49
bdl
0·62 0·21
0·20 0·11
bdl
101·50
91·6
OPX-L
5
57·98 0·23
bdl
2·41 0·13
0·82 0·10
0·15 0·14
4·38 0·08
35·74 0·19
0·57 0·05
bdl
bdl
102·04
93·6
41·64 0·63
bdl
bdl
bdl
0·36 0·19
8·89 1·09
50·29 1·08
bdl
bdl
bdl
101·18
91·0
41·74 0·11
bdl
bdl
bdl
0·31 0·21
9·26 0·15
49·84 0·11
bdl
bdl
bdl
101·16
90·6
0·16 0·11
6·12 0·29
34·40 0·48
0·78 0·12
bdl
bdl
101·37
90·9
4·49 0·21
35·47 0·41
0·63 0·05
bdl
bdl
101·38
93·4
91·6
Run P-9 (10008C, 1·45% H2O)
OL Lhz
5
OL-L
OPX Lhz
5
57·42 0·40
bdl
2·07 0·41
0·43 0·12
OPX-L
5
57·62 0·51
bdl
2·38 0·32
0·79 0·12
95% HZ 1þ 5% ’anhydrous phlogopite’ at 4·0 GPa
Run P-12 (11008C, 0·145% H2O)
OL Lhz
3
41·40 0·37
bdl
bdl
bdl
0·36 0·24
8·27 0·08
50·53 0·47
bdl
bdl
100·57
OPX Lhz
3
57·89 0·27
bdl
1·16 0·06
0·31 0·03
bdl
5·18 0·11
35·04 0·29
bdl
0·82 0·16
bdl
bdl
100·41
92·3
OPX-L
4
58·26 0·18
bdl
1·51 0·16
0·62 0·09
bdl
4·73 0·05
35·48 0·19
0·72 0·09
bdl
bdl
101·31
93·0
Run P-11 (11008C, 1·45% H2O); CPX** ¼ (50·11 wt%K2O)
OL Lhz
4
40·97 0·71
bdl
bdl
bdl
0·39 0·12
7·97 0·32
49·39 1·38
bdl
bdl
bdl
98·72
91·7
OL-L
2
41·34 0·01
bdl
bdl
bdl
0·51 0·00
8·49 0·12
49·51 0·07
bdl
0·18 0·18
0·12 0·07
100·14
91·2
OPX Lhz
6
55·97 1·03
bdl
1·28 0·49
0·40 0·26
bdl
5·19 0·32
33·53 0·85
0·84 0·13
0·32 0·38
0·23 0·26
97·75
92·0
OPX-L
5
57·42 0·89
bdl
1·61 0·50
0·62 0·24
bdl
4·45 0·43
35·03 0·49
0·67 0·10
bdl
bdl
99·80
93·4
91·8
Run P-8 (11508C, 1·45% H2O)
OL Lhz
6
41·14 0·25
bdl
bdl
bdl
0·34 0·15
7·96 0·20
49·77 0·31
bdl
bdl
bdl
99·22
OL-L
2
41·93 0·08
bdl
bdl
bdl
0·37 0·11
8·26 0·34
50·84 0·24
bdl
bdl
bdl
101·39
91·6
OPX Lhz
4
57·51 0·54
bdl
1·48 0·11
0·31 0·09
0·14 0·13
5·11 0·13
34·73 0·17
1·00 0·08
bdl
bdl
100·27
92·4
10
57·03 0·48
bdl
1·85 0·38
0·63 0·21
bdl
4·65 0·31
34·56 0·44
0·78 0·10
bdl
bdl
99·49
93·0
OPX-L
Run P-32 (11758C, dry)
Ol Lhz
4
41·13 0·37
bdl
bdl
bdl
0·44 0·13
8·29 0·29
49·74 0·36
bdl
bdl
bdl
99·59
91·5
OL-L
1
41·25
bdl
bdl
bdl
0·53
8·96
49·62
bdl
bdl
bdl
100·36
90·8
OPX Lhz
6
57·10 0·67
bdl
1·66 0·22
0·25 0·10
0·20 0·13
5·54 0·38
33·90 0·52
1·10 0·11
bdl
bdl
99·75
91·6
OPX-L
3
56·91 0·40
bdl
1·69 0·08
0·43 0·05
bdl
5·11 0·07
33·88 0·18
1·11 0·17
0·16 0·09
0·13 0·20
99·42
92·2
Three experiments [O60, O61, O89; Green et al. (2010, supplementary tables 3 & 4 therein)] used olivine discs inserted in
the capsule and surrounded by HZ2 Lherzolite mix. The three discs were cut from the same olivine crystal (San Carlos
olivine) and the reaction between the HZ2 Lherzolite mix and the olivine disc was examined in O89. Olivine analyses
located 410 microns from the margin of the disc or from a fracture decorated with clinopyroxene (Fig. S1f of Green et al.,
2010) have Mg# 91·5. Olivine within the lherzolite mix averages Mg# ¼ 93·9 and olivine within 10 microns of the margin
or of the clinopyroxene within the fracture averages Mg# ¼ 93·9. Our data show that the olivine disc composition (Mg#)
relevant to the FTIR measurements is that of the original San Carlos olivine and that only the margins (to approx 10
microns) of the olivine has re-equilibrated in Mg# with the more magnesian lherzolite HZ2. The data for O89 are tabulated
under O60 as no additional internal disc compositions were determined in this experiment. In the remainder of the
experiments with HZ2 Lherzolite and olivine layers, the observation that the starting material included many grains 520
microns diameter and samples clearly showed grain growth is consistent with olivine-in-layer compositions of Mg#¼ 91·5
(O53); 92·4 (O76); 92·6 (O52); 93·2 (O81) i.e. transitional between initial composition of Mg# ¼ 90–91·5 and HZ2 Lherzolite
composition Mg# 94. For experiments with HZ1 composition the San Carlos olivine composition of Mg# ¼ 90–91·5 is
very close to that within the lherzolite layer and this is apparent in this Table. Lhz ¼ lherzolite layer, -L ¼ sensor mineral
layer, disc ¼ sensor discs, bdl ¼ below detection limit
2077
JOURNAL OF PETROLOGY
VOLUME 53
Phase relations and chemical compositions
Experimental conditions and phase relations (Table 2) are
from Green et al. (2010). In the experiments with San Carlos
olivine layers (configuration 1) olivine, orthopyroxene,
clinopyroxene, garnet, pargasite and vapour were present
in the lherzolite below the solidus at 2·5 GPa, with pargasite
and aqueous vapour replaced by melt ( vapour) above the
solidus (Table 2). Phase relations at 4 GPa were similar
except that pargasite is missing from the assemblage (Table
2). It is important that addition of olivine to the lherzolite
composition does not alter the phase assemblage (other
than increasing the modal olivine content in the bulk
charge) except to modify the Mg# of phases in the HZ2
composition towards lower values. For the HZ1 and
HZ1 þPhlogopite compositions with Mg# 90, the addition of layers of olivine (Mg# ¼ 90) has no effect.
The experiments with Al-rich orthopyroxene and clinopyroxene layers (configuration 2) were conducted below
the solidus and with 1·45 wt % H2O. They produced olivine, orthopyroxene, clinopyroxene and garnet present
with vapour in the lherzolite layer at both 2·5 and 4 GPa.
Pargasite is absent at 2·5 GPa because the increased
modal clinopyroxene in the charge, owing to the clinopyroxene layer, lowers the Na2O content of clinopyroxene
(i.e. the activity of Na2O in the charge) below that necessary for pargasite stability (Table 2). In the mineral layers
the clinopyroxene grains are 30^60 mm in size, whereas
the orthopyroxene is 50^150 mm after the experiments.
The Al-rich orthopyroxene layer contained 6·3 wt % of
Al2O3 before the experiments, but this decreased to mean
compositions of 1·4 and 2·8 wt % at 4 and 2·5 GPa respectively, as Al2O3 exsolved to form garnet in the mineral
layer (Table 5). Similarly, the Al-rich clinopyroxene lost
most of its original Al2O3 (6·8 wt %) with mean compositions of 2·0 and 3·7 wt % at 4 and 2·5 GPa, respectively
(Table 5). The recrystallized pyroxenes are not homogeneous, as relict Al-rich cores were found. The composition
of orthopyroxene in the originally monomineralic layers is
on average more aluminous and has greater variability
than that in the lherzolite layers (Table 5). The decrease in
alumina is a consequence of exsolution of garnet and is
more marked at higher pressures.
In the experiments with Al-poor pyroxenes (configuration 3), the orthopyroxene lost some of its Al2O3 to form
garnet at 4 GPa, but there is little change at 2·5 GPa
(Table 5). Al-poor clinopyroxene has similar Al2O3 content
(2·6 wt %) at both pressures (Table 5). The resulting
compositions are more homogeneous than those from the
Al-rich pyroxenes of configuration 2, with minor variations only for Na2O and Al2O3 in the Al-poor clinopyroxene at 2·5 GPa.
Comparisons of clinopyroxene and orthopyroxene analyses between experiments with olivine layers and those
with pyroxene layers show that there is incomplete
NUMBER 10
OCTOBER 2012
equilibration of the pyroxenes throughout the charge, but
as already pointed out, there is sufficient exchange of
major elements between the lherzolite mix and the sensor
layers at 2·5 GPa to decrease the Na content of clinopyroxene in the lherzolite layer and consequently to destabilize
pargasite. The direction of reaction in configuration 2
(high-alumina pyroxenes) is to strongly decrease the Al
solubility in pyroxene by formation of garnet, and also to
decrease the Ca content of orthopyroxene and increase
the Ca content and Na content of the high-alumina clinopyroxene starting material. Collectively, the data show
that the experiments using low-alumina pyroxene layers
(configuration 3) or olivine þ low-alumina orthopyroxene
(configuration 4) yield products closest to the pyroxene
compositions in experiments with HZ1 or HZ1 þ
Phlogopite lherzolite, in which there were olivine-only
layers (configuration 1) or no layers at all. In interpreting
the phase relations, particularly pargasite stability and solidus temperature, Green et al. (2010, 2011) emphasized that
only those experiments without monomineralic layers or
with olivine only (configuration 1) were used to define the
phase relations and solidus temperatures of HZ1 and HZ2
lherzolites. This is because of the participation of the pyroxene layers in reaction with the lherzolite. With respect
to interpreting the FTIR data and water solubility in pyroxenes and olivine, the spectra obtained are specific to the
phases as analysed. Because the OH solubility in orthopyroxene and clinopyroxene is positively correlated with
AlAlMg1Si1 (Tschermak’s substitution; e.g. Stalder,
2004), the presence of relict alumina-rich cores in P4 and
P3 (at 2·5 and 4 GPa respectively) will bias water contents
to higher values than appropriate for the equilibrated lherzolite composition (Table 2). Thus experiments using configuration 3 (P6, P7) or configuration 4 (P8^P12, P32) are
preferred for deducing the contents of the lherzolite assemblage at the P, Tof the experiments. The experiments with
San Carlos olivine and Al-poor orthopyroxene layers (configuration 4) (i.e. HZ1 þ5 wt % ‘dry’ phlogopite) and
with 1·45 wt % H2O have olivine, orthopyroxene, clinopyroxene, garnet, phlogopite and vapor present at 4 GPa
and 11008C (P11) (i.e. below the solidus), but above the solidus quenched melt is present and phlogopite absent at
11508C (P8, Table 2). The P12 experiment at 4 GPa and
11008C with 0·145 wt % H2O has the same solid phases as
in P11 (1·45 wt % H2O) but mass-balance calculations
show that vapor is absent as all water is taken up by
NAMs and phlogopite. At 2·5 GPa and 10008C (P9 and
P10) pargasite is present in addition to the phlogopite and
garnet lherzolite assemblage, and vapor is present with
1·45 wt % H2O (P9) but absent with 0·145 wt % H2O
(from mass-balance calculations, P10, Table 2). The olivine
and orthopyroxene grains were 50^200 mm in the layers,
and free of inclusions and impurities. Olivine has an
Mg# of 90·6 at 2·5 GPa (P9, P10) and 91·6 at 4 GPa
2078
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
(P8, P11, P12) and is homogeneous (Table 5). Orthopyroxene is slightly inhomogeneous in layers at 4 GPa, as
Al2O3-rich relicts remained (Al2O3 42 wt %) after the
original orthopyroxene (2·9 wt %), but the newly crystallized orthopyroxene has an Al2O3 content around 1·3^
1·5 wt % (Table 5). At 2·5 GPa orthopyroxene composition
was homogeneous (2·4 wt % Al2O3).
The nominally ‘dry’ experiments of configuration 4 (i.e.
P30 and P32, Table 2) do contain a small amount of water
and show similar phase relations to other water-bearing
experiments, as, in addition to the NAMs, pargasite is present at 2·5 GPa in the HZ1 bulk composition (P30), as is
phlogopite at 4 GPa in the HZ1 þ5 wt % ‘dry’ phlogopite
bulk composition (P32). Vapour, however, is absent in
both experiments. The chemical compositions of olivine
and orthopyroxene in the ‘dry’ experiments closely resemble their counterparts in the experiments with 0·145 wt %
H2O (i.e. P10 and P12; Table 5).
Infrared spectroscopy
Olivine
The olivine in the two experiments that used discs of San
Carlos olivine (O60, O61; Table 3) displays a very broad
band of variable intensity around 3450 cm1, on which
some smaller structurally bound water bands are superimposed (Fig. 4). Planes of fluid inclusions were observed on
polished surfaces and in thin sections. The intensity of the
broad band correlates with the abundance of the visually
observed fluid inclusions. The broad band is attributed to
molecular water in fluid inclusions, and its presence
makes it difficult to quantify the structurally bound water.
Nevertheless, it provides clear evidence that an aqueous
fluid was present at the run conditions.
In the other experiments, where layers of crushed San
Carlos olivine were used (configuration 1), the background
was relatively smooth, the structurally bound water signal
was strongandthebroad feature around 3450 cm1appeared
rarely (Fig.4).The spectra of the re-equilibrated olivine in experimental charges are different from the absorption characteristics of the original San Carlos olivine, which has only
two major absorption bands at 3572 and 3525 cm1 [Ti]
(Fig. 3). These two bands are also present in some of our experiments, although a band at 3612 cm1 appears and becomes more significant with increasing pressure (Fig. 4,
Table 3). There is also a small shoulder on the 3568 cm1
band at 3574 cm1 in experiments at 4 GPa (Fig. 4b). Besides
the [Ti] and [Si] bands, there are also [triv] bands at 3354
and 3329 cm1 (Fig. 4). The relative proportion of the latter
bands is larger in the experiment using Au25Pd75 capsules
(O98,Table 3), which is at 4 GPa and12258C.
At 2·5 GPa, in the experiments with San Carlos olivine
layers (configuration 1), the water concentration in olivine
is slightly different for each of the two lherzolite mixes
(i.e. HZ1 and HZ2), being higher in experiments with
the more iron-rich HZ1 mix (67^71ppm) (i.e. O85, O79,
O77) compared with HZ2 (31^52 ppm) (i.e. O52, O53)
(Tables 2 and 3). The contribution of the [triv] bands remains the same, irrespective of which lherzolite mix is
used for the experiment. Water concentrations are similar
at 2·5 GPa, within analytical uncertainty, whether the
system is above (O79, 71ppm; O77, 68 ppm) or below
(O85, 67 ppm) the solidus for the HZ1 composition
(Tables 2 and 3). A similar observation applies for the
HZ2 composition at 2·5 GPa: O53 (52 ppm) is above the
solidus and O60 (35 ppm) and O52 (31ppm) are below
the solidus. The effect of the initial water concentrations
could not be studied in the experiments with San Carlos
olivine layers at 2·5 GPa as we did not vary initial water
concentrations while keeping other variables constant.
At 4 GPa, where initial water concentration was varied
(0·145 and 1·45 wt %), the water concentrations in olivine
of configuration 1 (62^188 ppm) below the solidus in both
mixes (HZ1 and HZ2) (i.e. O76, O80, O81 and O92) are
generally higher than those at 2·5 GPa (67 ppm; i.e. O85)
(Tables 2 and 3). The olivine above the solidus at 4 GPa in
mix HZ1 (O98) has much less water (30 ppm) than subsolidus runs (110^188 ppm) (Table 2). In the subsolidus experiments at 4 GPa, 11008C with a water-rich vapour
phase (O80 and O92), the initial water concentration appears to play a role, as olivine has 188 ppm (O80) and
110 ppm (O92) of water for 1·45 and 0·145 wt % initial
water concentrations respectively (Table 2). It should be
noted that the relative proportions of the different substitutions are similar in these two experiments (Table 3 and
Fig. 4).
San Carlos olivine in configuration 4 shows similar
absorption characteristics to those in configuration 1, with
only moderate variation with pressure (Tables 2 and 3,
Fig. 6a). The contribution of the [Si] 3612 cm1 band
slightly increases with increasing pressure in both configuration 1 and 4 experiments (Fig. 6a, Table 3), and this tendency was also noticed by Mosenfelder et al. (2006) in
simple experimental systems (i.e. natural Fe-bearing olivine and synthetic forsterite) between 2 and 12 GPa. At subsolidus conditions, olivine in the K-enriched experiments
of configuration 4 with HZ1 þ5 wt % ‘dry’ phlogopite
mix tends to lower concentrations of water, 62^71ppm at
4 GPa and 44 ppm at 2·5 GPa, with respect to the counterpart olivine in configuration 1 with the HZ1 lherzolite mix
(110^188 ppm at 4 GPa and 67^71ppm at 2·5 GPa) (Tables
2 and 3, Figs 4 and 6).
At 4 GPa, San Carlos olivine in configuration 4 contains
somewhat more water than at 2·5 GPa (Table 2). However,
as for configuration 1, water concentrations are similar
below (P8) and above the solidus (P11) at 4 GPa within
analytical uncertainty (Table 2).
In configuration 4 at 2·5 and 4 GPa, the same
water-related absorption bands were identified in olivine
2079
(b)
1000 oC, 2.5 GPa, HZ2, 2.9 wt% 40
(O60)
1000 oC, 2.5 GPa, HZ1, 1.45 wt%
(O85)
1025 oC, 2.5 GPa, HZ2, 1.45 wt%
(O52)
1025 oC, 2.5 GPa, HZ1, 1.45 wt%
(O79)
NUMBER 10
OCTOBER 2012
40
1100 oC, 4 GPa, HZ2, 2.9 wt%
(O61)
1150 oC, 4 GPa, HZ2, 1.45 wt%
(O76)
Absorbance/cm
(a)
VOLUME 53
1100 oC, 4 GPa, HZ1, 1.45 wt%
(O80)
1100 oC, 4 GPa, HZ1, 0.145 wt%
(O92)
Absorbance/cm
JOURNAL OF PETROLOGY
1200 oC, 4 GPa, HZ1, 1.45 wt%
(O81)
1050 oC, 2.5 GPa, HZ2, 1.45 wt%
(O53)
1225 oC, 4 GPa, HZ1, 1.45 wt%
(O98/O99)
1050 oC, 2.5 GPa, HZ1, 1.45 wt%
(O77)
0
3700 3600 3500 3400 3300 3200 3100 3000
Wavenumber (cm--1)
0
3700 3600 3500 3400 3300 3200 3100 3000
Wavenumber (cm--1)
Fig. 4. Infrared spectra for olivine in ‘sensor’ layers of chemically complex, configuration 1 experiments at 2·5 GPa (a) and 4 GPa (b). The
increased solubility of water into olivine with pressure is demonstrated, whereas the water solubility does not vary with temperature below
and across the solidus of model mantle lherzolite. The spectra are stacked for clarity, and the dashed lines represent the background for each
spectrum.
(b)
60
Al-rich (2554)
1150 oC, 4 GPa, HZ1, 1.45 wt%
(P3)
60
50
50
Al-rich (2554)
1000 oC, 2.5 GPa, HZ1, 1.45 wt%
40
(P4)
Al-rich (2554)
1000 oC, 2.5 GPa, HZ1, 1.45 wt%
40
(P4)
Al-poor (2501)
1150 oC, 4 GPa, HZ1, 1.45 wt%
(P6)
30
Al-poor (2501)
1150 oC, 4 GPa, HZ1, 1.45 wt%
(P6)
20
30
Absorbance/cm
Al-rich (2554)
1150 oC, 4 GPa, HZ1, 1.45 wt%
(P3)
Absorbance/cm
(a)
20
Al-poor (2501)
Al-poor (2501)
1000 oC, 2.5 GPa, HZ1, 1.45 wt% 10
(P7)
o
1000 C,2.5 GPa, HZ1, 1.45 wt%
(P7)
10
0
3800 3700 3600 3500 3400 3300 3200 3100 3000
3800 3700 3600 3500 3400 3300 3200 3100 3000
Wavenumber (cm-1)
Wavenumber (cm-1)
Fig. 5. Infrared spectra for clinopyroxene (a) and orthopyroexene (b) in ‘sensor’ layers of chemically complex, configuration 2 and 3
experiments.
2080
(a)
WATER IN NOMINALLY ANHYDROUS MINERALS
(b)
1000 oC, 2.5 GPa, 1.45 wt% 20
(P9)
1000 oC, 2.5 GPa, 0.145 wt%
(P10)
10
1100 oC, 4 GPa, 1.45 wt%
(P11)
1100 oC, 4 GPa, 0.145 wt%
(P12)
1000 oC, 2.5 GPa, ‘dry’
(P30)
Absorbance/cm
1150 oC, 4 GPa, 1.45 wt%
(P8)
50
1000 oC, 2.5 GPa, 0.145 wt% 40
(P10)
15
1000 o C, 2.5 GPa, ‘dry’
(P30)
1000 oC, 2.5 GPa, 1.45 wt%
(P9)
30
1150 oC, 4 GPa, 1.45 wt%
(P8)
1100 oC, 4 GPa, 1.45 wt%
(P11)
5
1100 oC, 4 GPa, 0.145 wt%
(P12)
1100 oC, 4 GPa, ‘dry’
(P32)
20
Absorbance/cm
KOVA¤CS et al.
10
1100 oC, 4 GPa, ‘dry’
(P32)
3800 3700 3600 3500 3400 3300 3200 3100 3000
3800 3700 3600 3500 3400 3300 3200 3100 3000
Wavenumber (cm-1)
Wavenumber (cm-1)
Fig. 6. Infrared spectra for olivine (a) and orthopyroxene (b) in ‘sensor’ layers of chemically complex, configuration 4 experiments. The spectra
are stacked for clarity, and the dashed lines represent the background for each olivine spectrum.
in the nominally ‘dry’ experiments (P30, P32) as in the experiments with low water contents (i.e. 0·145 and 1·45 wt
% water) (Fig. 6). At 2·5 GPa and at subsolidus conditions,
the olivine in the ‘dry’ P30 experiment (HZ1 mix at
2·5 GPa, 10008C) has less water (32 ppm) than olivine in
configuration 1 experiments at the same P, T conditions,
and also with HZ1 mix but with 1·45 wt % water (67 ppm
water, O85, Tables 2 & 3). However, olivine in the ‘dry’
P30 HZ1 run has only slightly lower water within analytical uncertainty than olivine at the same P, Tconditions in
the HZ1 þ5 wt % ‘dry’ phlogopite mix of configuration 4
(P9 and P10) (44 ppm). At 4 GPa, the olivine in the subsolidus nominally ‘dry’ run P32 (HZ1 þ5 wt % ‘dry’ phlogopite mix, at 11758C) has a similar water concentration
(57 ppm) to its water-bearing subsolidus counterparts with
0·145^1·45 wt % water at 11008C (65^71ppm, P11 and P12),
and near-solidus run at 11508C of the same bulk composition (62 ppm, P8, Table 2).
Orthopyroxene
The spectra of orthopyroxene in configurations 2 and 3
with three major, broad bands at 3600, 3530 and
3420 cm1 all have similar absorption characteristics, regardless of the original Al2O3 content of pyroxenes as
well as the P^T conditions of the experiments (Fig. 5).
Such spectra are broadly similar to orthopyroxene from
natural mantle xenoliths from various tectonic settings
(Bell et al., 1995; Rossman, 1996; Peslier et al., 2002; Grant
et al., 2007a; Falus et al., 2008; Li et al., 2008; Yang et al.,
2008; Bonadiman et al., 2009; Sundvall, 2010; Xia et al.,
2010). Also, Rauch & Keppler (2002) and Stalder (2004)
observed similar bands in Al-doped enstatite in chemically
simple systems (i.e. MgO^Al2O3^SiO2^H2O).
At subsolidus conditions (configurations 2 and 3),
the initially Al-rich orthopyroxene always has more water
(472 and 377 ppm at 2·5 and 4 GPa, respectively) than the
initially Al-poor counterparts (365 and 260 ppm at 2·5
and 4 GPa, respectively) (Fig. 5, Tables 2 and 4). The difference is less significant at 2·5 GPa. This observation may be
a consequence of either increased incorporation of water
into more Al-rich pyroxene (i.e. Stalder, 2004; O’Leary
et al., 2010; and also see below), or incomplete equilibration
of the original pyroxene compositions, which can be identified and avoided in EDS analyses but cannot be excluded
in the larger areas analyzed by FTIR.
Orthopyroxene in both configurations 2 and 3 has more
water (and more Al2O3) at lower pressures; that is, the originally Al-rich and Al-poor orthopyroxenes (E2554 and
DTP1) both contain slightly more water (472 and 365 ppm
respectively) at 2·5 GPa than at 4 GPa (377 and 260 ppm
respectively, Fig. 5, Tables 2 and 4). The effect of pressure
is less significant for Al-rich orthopyroxene.
The orthopyroxene in configuration 4 has three major
broad bands at 3600, 3530 and 3420 cm1 (Fig. 6), which
are identical to the other experiments with a different arrangement of the orthopyroxene sensor layers (i.e. configurations 2 and 3) (Fig. 5). In addition, orthopyroxene
of configuration 4 in the nominally ‘dry’ experiments
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JOURNAL OF PETROLOGY
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(P30, P32) shows the same water-related absorption bands
as the counterparts with low water contents (configurations 3 and 4; Fig. 6). At 4 GPa and at subsolidus conditions, the water concentrations in orthopyroxene appear
to be independent of the initial water concentrations at
otherwise identical conditions (i.e. P12, P11 and P32 with
0·145 and 1·45 wt % water, and nominally ‘dry’, respectively) (Tables 2 and 4). The experiment at 4 GPa and
11508C (P8) above the solidus has similar water concentration in orthopyroxene (311ppm) to its counterpart below
the solidus at 11008C (P11, 317 ppm) both with 1·45 wt %
bulk water (Tables 2 and 4). These concentrations are the
same within analytical uncertainty as those in Al-rich
orthopyroxene in configuration 2 and 3 experiments
(260 ppm) at 4 GPa (Table 2).
Clinopyroxene
The re-equilibrated clinopyroxene in configurations 2 and
3 has two broad bands at 3640 and 3450 cm1 and a smaller one at 3360 cm1 (Fig. 5). This differs from the two original clinopyroxenes both in the position (Fig. 3) and the
considerably higher intensities of the absorption bands
(Fig. 5). These spectra closely resemble some other natural,
mantle-derived and experimentally synthesized diopsidic
clinopyroxene (Skogby, 1994; Ingrin et al., 1995; Peslier
et al., 2002; Andrut et al., 2003; Bell et al., 2004; Bromiley
et al., 2004; Aubaud et al., 2007; Grant et al., 2007b; Stalder
& Ludwig, 2007; Li et al., 2008; Yang et al., 2008;
Bonadiman et al., 2009; Sundvall, 2010; Xia et al., 2010;
Fig. 5). In the experiments with double pyroxene layers
(configurations 2 and 3), the initially more Al-rich clinopyroxene always has more water [1590 ppm at 4 GPa (P3)
and 1431ppm at 2·5 GPa (P4)] than their initially Al-poor
counterparts [910 ppm at 4 GPa (P6) and 978 at 2·5 GPa
(P7)], and clinopyroxene has always more water than the
coexisting orthopyroxene (Figs 5 and 6, Tables 2 and 4).
Lherzolite matrix
The fine-grained lherzolite layers were also analysed by
FTIR to determine whether the presence or absence of
hydrous phases (pargasite and phlogopite) could be detected. This is possible because pargasite and phlogopite
have strong, distinctive and non-overlapping absorption
bands above 3650 cm1, and therefore can be distinguished
from both structurally bound water in NAMs and water
vapour. Pargasite has a number of major bands at
3715^3650 cm1 (Della Ventura et al., 2003, 2007), whereas
phlogopite has fewer, but more pronounced bands at
higher wavenumbers (3725^3670 cm1; Chaussid, 1970;
Wunder & Melzer, 2002).
In experiments O80, O79, P3, P6 and P8 where hydrous
phases were not detected by SEM, there is no indication
of any pargasite- or phlogopite-related bands in the IR
spectra (Fig. 7). In nominally ‘dry’ P30 and O85 with
1·45 wt % bulk water, where pargasite is the only stable
NUMBER 10
OCTOBER 2012
hydrous phase in the lherzolite matrix, there are two
major bands at 3711 and 3673 cm1 with two smaller ones
at 3689 and 3650 cm1 (Fig. 7). Where phlogopite is the
stable hydrous phase (i.e. nominally ‘dry’ P32 and P11
with 1·45 wt % H2O; Table 2) we observe three major
bands at 3722, 3717 and 3712 cm1 (Fig. 7). There are some
instances at 2·5 GPa (i.e. P9, P10, Table 2) where both hydrous minerals are stable and the spectra show both pargasite and phlogopite (Fig. 7). Thus, the two nominally ‘dry’
experiments of configuration 4 at 2·5 and 4 GPa, P30 and
P32, respectively, crystallized pargasite and phlogopite
(P30), or phlogopite (P32) in the lherzolite layer. It is
clear that the experimental techniques did not entirely exclude water in these ‘dry’ experiments. Pargasite was not
detected by either SEM or FTIR in configurations 2 and
3, and this is attributed to the bulk compositions of these
layered experiments where higher modal clinopyroxene
accommodates Na, Al and Ti and destabilizes the pargasite
crystallized in HZ1 or HZ2 lherzolite þ olivine (configuration 1), or HZ1 lherzolite þ olivine þ orthopyroxene or
HZ1 þ5 wt % ‘dry’ phlogopite þ olivine þ orthopyroxene
(configuration 4) compositions (Green et al., 2010). Thus,
the presence of vapour in runs of configurations 2 and 3
at 2·5 GPa (P4 and P7) is a result of excess water, which
cannot be accommodated in HZ1 þpyroxene bulk
compositions.
DISCUSSION
Comparison with other approaches
The question of water incorporation in upper mantle
minerals has previously been addressed using different
analytical protocols, which have dictated different experimental approaches. The non-destructive analysis by
IR-spectroscopy has the advantage of providing information about the substitution mechanism of water in NAMs
(e.g. Berry et al., 2005), and of quantifying various
H-species such as H2O and OH (i.e. Hauri et al., 2002).
However, quantification of water in NAMs by using
IR-spectroscopy was compromised because it had been
thought that quantitative results could only be obtained
using polarized light and three measurements along mutually orthogonal sections (Paterson, 1982; Libowitzky &
Rossman, 1996, 1997). The production of crystals large
enough for this approach cannot be routinely achieved in
phase-equilibrium experiments. To circumvent this obstacle two approaches have been employed within the experimental community.
One approach was to run large cubes of sensor crystals
in simple chemical systems using excess water to guarantee
satisfactory grain size (e.g. Rauch & Keppler, 2002; Hauri
et al., 2006; Aubaud et al., 2007; Bali et al., 2008). There are
several disadvantages to this method. First, the sensor crystals usually have their own, inherited, point-defect concentrations and trace-element chemistry, which may not be
2082
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
120
Phlogopite+
pargasite, P10
(4) 1000 oC, 2.5 GPa, 0.145 wt% 100
Phlogopite, P32
(4) 1175 oC, 4 GPa, ‘dry’
Phlogopite, P11
80
Pargasite, P30
(4) 1000 oC, 2.5 GPa, ‘dry’
Pargasite, O85
(1) 1000 oC, 2.5 GPa, 1.45 wt%
60
No hydrous phase, P8
Absorbance/cm
(4) 1100 oC, 4 GPa, 1.45 wt%
(4) 1150 oC, 4 GPa, 1.45 wt%
40
No hydrous phase, P6
(3) 1150 oC, 4 GPa, 1.45 wt%
No hydrous phase, P3
(2) 1150 oC, 4 GPa, 1.45 wt%
No hydrous phase, O79
No hydrous phase, O80
3800
3750
3700
(1) 1025 oC, 2.5 GPa, 1.45 wt%
20
(1) 1100 oC, 4 GPa, 1.45 wt%
3650
3600
3550
0
3500
Wavenumber (cm--1 )
Fig. 7. Infrared spectra for the fertile lherzolite matrix. Continuous
lines indicate major bands for pargasite and phlogopite, where these
phases are present.
equilibrated during the experiments. The post-experiment
water contents may then represent metastable conditions.
For this reason we examined carefully the sensor minerals
before and after the experiments (Fig. 3) to demonstrate
that new bands corresponding to the new equilibrium conditions are formed. Some bands remained from the starting minerals, which does not necessarily mean that
equilibration is incomplete, as it could well be that these
bands (and the substitution mechanisms they represent)
are part of the equilibrium under the new conditions.
Second, the point defects so important for water substitution in NAMs may not be those appropriate for natural
upper mantle mineral assemblages such as peridotite in
which the chemical potentials of major components are
buffered by the phase assemblage (e.g. silica activity by
olivine þ orthopyroxene). Third, minor elements such as
Ti are generally not included, although it has been shown
that Ti is essential for understanding the water incorporation into upper mantle olivine (Berry et al., 2005; Walker
et al., 2007). As pointed out by Walker et al. (2007), the stoichiometries of the four main water substitution mechanisms in olivine are different, with the possibility that for
the [Si] mechanism, C[Si]H2O / (fH2O)2, for [Mg],
C[Mg]H2O / fH2O, whereas for [triv] and [Ti], C[triv &
1/2
Ti]H2O / (fH2O) . One consequence is that the [Si]
and [Mg] mechanisms might predominate at high fH2O
but be less important than [Ti] and [triv] at waterundersaturated conditions at identical temperature and
pressure. The prevalence of [Ti] and [triv] is indeed a feature of this study. Although the effect of varying fH2O at
constant Tand P is yet to be tested experimentally, the possibility of different substitution stoichiometries in olivine
invalidates current extrapolation of simple system results
obtained at very high relative fH2O to realistic melting
scenarios in natural upper-mantle compositions. Experiments in simple systems generally do not access the T^P^
fH2O conditions relevant to fertile peridotite because of
the much lower solidus temperature of peridotite at given
P and fH2O in comparison with the simple systems
(i.e. Grant et al., 2006; Aubaud et al., 2007). In addition,
the lack of elements such as Na and K that readily enter
the fluid phase often means that fH2O is higher in simple
systems at vapour saturation at given Tand P.
Another experimental approach to achieve sufficient
grain size to investigate water incorporation in NAMs is
to start the experiments several hundred degrees above
the target temperature, to allow formation of large crystals
upon cooling to the target temperature (i.e. Grant et al.,
2006, 2007b; Aubaud et al., 2007). However, the water defects may be characteristic of the higher temperature
stage if diffusion is not rapid enough to establish a new
equilibrium at lower temperatures. Also, although diffusion of water in clinopyroxene may take place very rapidly
(e.g. Hercule & Ingrin, 1999), complete compositional
equilibration of the higher temperature mineral phases
during subsequent cooling is rarely if ever achieved in
such experiments, leaving compositional core-to-rim
variations.
Alternatively, the water in NAMs from experiments has
been determined by SIMS (e.g. Hauri et al., 2002; Koga
et al., 2003; Aubaud et al., 2007, 2009; Tenner et al., 2009;
O’Leary et al., 2010). SIMS is able to analyse smaller
grains than those suitable for polarized IR spectroscopy
and thus is applicable for a much wider range of experiments. Nevertheless, synthesized minerals need to be at
least 20 mm in size for SIMS (i.e. Hauri et al., 2002). As a
consequence, this technique has been mainly applied to
melting experiments where minerals with sufficient grain
sizes coexist with a hydrous melt (Aubaud et al., 2004;
Hauri et al., 2006; Tenner et al., 2009; O’Leary et al., 2010),
whereas experimental subsolidus or near-solidus mineral
phases are often too fine-grained to be measured.
Although this approach provides important information
on water partitioning at supra-solidus temperatures, very
little information is available on upper mantle lherzolitic
subsolidus and near-solidus conditions, or at varying fH2O.
In addition, the melt þ crystal(s) approach obtains data
for the particular mineral and melt compositions but consideration of hydrous melting of lherzolite requires knowledge of the equilibrium compositions of all phases, and of
melt in equilibrium with them. For example, clinopyroxene
coexisting with olivine on the liquidus of olivine basalt at
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JOURNAL OF PETROLOGY
VOLUME 53
2 GPa differs in composition from that which is in equilibrium with olivine þ orthopyroxene þ clinopyroxene þ garnet þ melt at the same P^T. The method used here
ensures that composition of the NAMs of interest are
those appropriate to the fertile lherzolite solidus.
The similarity of absorption characteristics in our experimental runs to those in mantle-derived rocks indicates
that hydrogen defects in NAMs can be successfully reproduced in chemically complex experiments using the
layered experimental technique. We have demonstrated
that most substitution mechanisms are not inherited from
the original sensor crystals. It is thus possible for the first
time to match the water contents and hydrogen defects in
NAMs with the mineralogy (particularly the presence of
hydrous phases) and melting relations of natural mantle
lherzolite compositions at uppermost mantle conditions.
Water storage capacity of NAMs in the
lherzolitic mantle
If we take the concentrations from NAMs in our representative configuration 3 and 4 experiments (Tables 2 and 6)
and combine them with the estimated modal composition
of garnet lherzolitic mantle used by Hirth & Kohlstedt
(1996) (i.e. olivine 56%, orthopyroxene 19%, clinopyroxene 10%, garnet 15%), we find that the storage capacity of
the peridotitic upper mantle at pressures of 2·5 and 4 GPa
is 185 and 197 ppm respectively (hereafter we will use the
mean value of 190 ppm; Table 6).
We take configuration 3 and 4 experiments for this estimation (P6, P7, P8, P11 and P12, Table 2) because these
are the ones where Al-poor pyroxenes and olivine coexist
in the fertile lherzolite (HZ1, HZ1 5 wt % ‘dry’ phlogopite) mix. As noted previously, Al-poor pyroxenes show a
much better degree of equilibration (i.e. do not preserve
Al-rich cores during re-equilibration unlike their Al-rich
counterparts), therefore closely approaching equilibrium
conditions. In our calculation we assume that the garnet
has a similar amount of water to olivine (Ingrin &
Skogby, 2000). There are only a very limited number of experimental data on the solubility of water in garnet, but
these generally show similar or slightly higher water concentration than in olivine, especially at pressures 44 GPa
(Lu & Keppler, 1997; Withers et al., 1998; Hauri et al.,
2006). The uncertainty in partitioning of water between
olivine and garnet does not introduce a major error in our
calculation, as even if there was an order of magnitude
more water in garnet than olivine it would translate only
to an 10% increase in bulk water stored in peridotitic
NAMs at uppermost mantle conditions as the estimated
modal abundances of peridotitic mineral phase assemblages remain approximately the same over the P, Tconditions studied. Towards greater mantle depths, the modal
abundances of garnet and pyroxenes change significantly,
with implications for the storage capacity (i.e. Tenner
et al., 2012). The water budget of peridotite is dominated
NUMBER 10
OCTOBER 2012
by the pyroxenes. Although clinopyroxene makes up only
10% the NAMs of fertile peridotite, it accounts for more
than 50% of their water storage capacity. Therefore, the
storage capacity of a garnet harzburgite is significantly
less than that of a fertile garnet lherzolite at the same P
and T. The estimate of the maximum water storage capacity in NAMs (190 ppm) in lherzolite is in agreement
with the water concentration in the peridotitic upper
mantle source for mid-ocean ridge basalts (MORB) of
50^200 ppm (Dixon et al., 1988; Michael, 1988, 1995;
Danyushevsky et al., 2000; Saal et al., 2002). The maximum
water storage in upper mantle NAMs, however, falls considerably below estimates for ocean island basalt (OIB) or
E-MORB (‘enriched’) sources of 300^1000 ppm (Dixon
et al., 2002; Hauri et al., 2002; Asimow et al., 2004; Tenner
et al., 2012; and references therein).
If the bulk water content is higher than the 190 ppm
peridotite storage capacity then pargasite, phlogopite
(if K2O is sufficient) and/or a water-rich vapour phase
will be present at subsolidus conditions with hydrous
silicate melt forming above the dehydration solidus or
vapour-saturated solidus (Green et al., 2010). In the uppermost peridotitic mantle (i.e. 53 GPa, equivalent to a
mantle depth of 90 km), pargasite is the major subsolidus
host for any water in excess of the saturation limit of the
NAMs. The composition of pargasite in lherzolite varies
as a function of P and T (Niida & Green, 1999).
Accordingly, the maximum water storage capacity of
pargasite peridotite may vary from 0·5 wt % (30%
pargasite þ NAMs) at 1·5 GPa, 10008C to 0·1wt %
(5% pargasite þ NAMs) at 3 GPa, 10008C, assuming
1·5 wt % H2O in pargasite. These estimations are based
on the experimental study of Niida & Green (1999), where
the modal abundance of pargasite at different pressures
was calculated from the chemical composition of the mineral phases by applying the method of Le Maitre (1979).
The actual storage capacity may be lower than this if the
modal proportion of pargasite is lower because of insufficient amounts of minor-element components needed to stabilize pargasite (i.e. Na2O, K2O, TiO2). Phlogopite may
increase the water storage capacity of peridotite in compositions with sufficient K2O to stabilize this phase. It
should be noted that the water residing in residual NAMs
(190 ppm at the vapour-saturated solidus) is a function
of temperature above the solidus, as the water content decreases in the increasing melt fraction and partition coefficients change with temperature and phase compositions.
A comprehensive review of published concentrations of
water in the peridotitic NAMs of the upper mantle determined by either FTIR using the mineral-specific calibrations of Bell et al. (1995, 2003) or SIMS using the
calibration of Koga et al. (2003) and Aubaud et al. (2007)
is summarized in Table 6 and compared with our results
in Fig. 8. The NAMs in our experiments display higher
2084
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
Table 6: Estimate for the bulk effective water capacity of lherzolitic upper mantle based on configuration 4 experiments at
2·5 and 4 GPa and comparison with other experimental studies using natural analogue starting composition and data from
natural mantle xenoliths
Reference:
This study1
Aubaud et al. (2004)
Hauri et al. (2006)
Tenner et al. (2009)
System:
experiments
experiments
experiments
experiments
Pressure (kbar):
40
25
10–20
10–40
30–50
Temperature (8C):
1100
1000
1230–1380
1000–1380
1350–1440
Analytical method:
FTIR
FTIR
SIMS
SIMS
SIMS
Calibration:2
B
Ko
Ko
A
K3
K3
B
min.
H2O ppm in ol
max.
min.
max.
min.
max.
42
66
30
44
64
139
36
94
29
H2O ppm in opx
309
309
295
295
692
1202
233
1840
436
996
H2O ppm in cpx
910
910
978
978
804
1587
332
2000
794
1426
Dopx/ol
4
Dcpx/opx4
bulk lherzolite5
7·3
4·7
9·9
6·7
7·9
2·94
2·94
3·32
3·32
1·38
180
197
175
185
9·7
257
486
6·5
15·5
0·9
1·4
103
25·39
1·07
616
215
1·23
371
Mantle xenoliths (spinel and garnet lherzolites, xenocrysts)
Reference:
Bell & Rossman (1992)
min.
H2O ppm in ol
max.
Peslier et al. (2002);
Grant et al.
Li et al.
Yang et al.
Falus et al.
Bonadiman
Sundvall
Xia et al.
Peslier & Luhr (2006)
(2007a)
(2008)
(2008)
(2008)
et al. (2009)
(2010)
(2010)
min.
min.
min.
min. max. min. max.
max.
max.
min.
max.
min.
max.
min.
max.
max.
1
79
0
7
3
53
2
45
0
35
2
15
—
—
—
—
0
0
H2O ppm in opx
50
460
39
272
169
201
53
402
5
140
92
305
9
92
10
281
8
94
H2O ppm in cpx
150
1080
140
528
342
413
171
957
5
355
186
632
5
399
59
629
27
223
6
60
18
355
3
68
7
41
3
8
18
81
Dopx/ol
4
Dcpx/opx
4
bulk lherzolite5
1·50
38
3·00
132
1·30
21
2·90
109
1·90
68
2·30
117
1·30
29
3·4 0·75
204
1
3·4
87
1·70
38
2·60
132
0·50
6·50
1·9
6
6·5
56
1
The given values are averages of values reported in Table 2. For 4 GPa ol and opx concentrations are from P8, P11 and
P12; for 2·5 GPa ol and opx concentrations are from P9 and P10. Cpx concentrations are from P6 and P7 for 4 and 2
·5 GPa respectively.
2
B, Bell et al. (1995) for cpx and opx; K, Kovács et al. (2010); Ko, Koga et al. (2003); A, Aubaud et al. (2007).
3
cpx and opx concentrations are the same as for the Bell et al. (1995) calibration.
4
The partitioning coefficients from the cited references are recalculated from the water concentrations in coexsisting
NAMs, and as such may be slightly different from those reported in the references.
5
For the garnet lherzolite calculation it is assumed that ol and grt have similar concentration of H2O (Bell & Rossman,
1992; Ingrin & Skogby, 2000; Grant et al., 2007a). Water concentration data for NAMS are taken from Table 2.
Modal composition of garnet lherzolite is taken from Hirth & Kohlstedt (1996) [olivine (ol) 56%; clinopyroxene (cpx) 10%;
orthopyroxene (opx) 19%; garnet (grt) 15%].
water concentrations than most of the mantle peridotites,
with overlap only with the most water-rich natural samples
from supra-subduction environments (Bell & Rossman,
1992; Falus et al., 2008; Li et al., 2008; Table 6; Fig. 8). This
is consistent with our experiments reproducing the maximum water contents of the upper mantle in chemically
realistic complex systems where all the buffering phases,
olivine, clinopyroxene, orthopyroxene, garnet and/or pargasite and/or phlogopite, are simultaneously present. The
fact that most natural samples (especially those from cratonic areas and within-plate settings) show lower concentrations than our experiments may indicate that water
contents in the mantle distant from active subduction are
below the mantle’s effective storage capacity in NAMs.
The water concentrations in NAMs in our study are,
nevertheless, generally lower than what has been reported
from other experiments, where water in NAMs was measured by SIMS (Aubaud et al., 2004; Hauri et al., 2006;
2085
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
[this study], n=5
[1], n=9
This study (experimental)
[2], n=10
[3], n=1
[4], n=6
[5], n=17
[6], n=8
[7], n=17
[8], n=27
[9], n=8
[11], n=46
Olivine
Experimental
Mantle xenoliths and xenocrysts
[this study], n=5
[1], n=2
[2], n=14
[3], n=6
[4], n=7
[5], n=15
[6], n=8
[7], n=13
[8], n=26
[9], n=8
[10], n=46
[11], n=46
[12], n=8
Orthopyroxene
[this study], n=5
[1], n=2
[2], n=14
[3], n=6
[4], n=7
[5], n=15
[6], n=8
[7], n=13
[8], n=26
[9], n=8
Clinopyroxene
[10], n=46
[11], n=35
[12], n=14
100
200
300
400
500
600
700
800
900
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
H2O (ppm)
Fig. 8. Water concentrations (ppm wt % H2O) in various NAMs from natural upper mantle peridotite xenoliths and experiments. References:
[1], Aubaud et al. (2004); [2], Hauri et al. (2006); [3], Tenner et al. (2009); [4], Bell & Rossman (1992); [5], Peslier et al. (2002) and Peslier & Luhr
(2006); [6], Grant et al. (2007a); [7], Li et al. (2008); [8], Yang et al. (2008); [9], Falus et al. (2008); [10], Bonadiman et al. (2009); [11], Xi et al. (2010);
[12], Sundvall (2010). n, number of samples used for the plot of the particular range. The ranges for particular NAMs from this study are plotted
based on P8^P12 for olivine and orthopyroxene and P6 and P7 for clinopyroxene in Table 2.
Tenner et al., 2009; Table 6). This difference is related to the
different experimental strategies, as these latter experiments were undertaken in basaltic systems with high degrees of thermodynamic variance. Under such conditions,
the chemical potentials of the major components may
differ from those in mantle assemblages, stabilizing different
defect mechanisms for H2O substitution in NAMs.
Similarly, the high initial water concentration (45 wt %)
may also result in different H2O substitution mechanisms.
In principle, experiments conducted at high relative fH2O
2086
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
cannot be extrapolated to the low relative fH2O likely to be
typical of most of the Earth’s mantle without establishing
the relationship between H2O concentration and fH2O. In
olivine, for example, this relationship is extremely complex
because of the existence of four substitution mechanisms
for H2O, each with different stoichiometries and hence different dependences on fH2O (Berry et al., 2005, 2007; Walker
et al., 2007; Kova¤cs et al., 2010; Balan et al., 2011).
Partitioning of water between NAMs and
hydrous phases.
The partition coefficient between clinopyroxene and
orthopyroxene (Dcpx/opx) in our representative experiments (P6 and P7 with Al-poor starting pyroxenes) is
3·5 1·5 and 2·7 1·1 at 4 and 2·5 GPa respectively
(Table 2). The average partitioning between coexisting
orthopyroxene and olivine (Dopx/ol) in our representative
experiments (P8^P12) is 4·7 1·1 and 6·7 2 at 4 and
2·5 GPa respectively (Table 2). These values are determined by using the average water concentrations in olivine
(66 and 44 ppm at 4 and 2·5 GPa respectively) and orthopyroxene (309 and 295 ppm at 4 and 2·5 GPa respectively)
from experiments P8^P12 (Table 2). These partition coefficients are within the large range of values determined for
natural mantle peridotites; that is, 3^355 and 0·5^6·5 for
Dopx/ol and Dcpx/opx respectively (Table 6 and Fig. 9 and
references therein).
The experimental studies, including ours, however, seem
to show generally lower Dopx/ol values than natural samples
(3-355). The reason for this discrepancy may be related to
preferential water loss from olivine during heating of the
mantle xenoliths by their host magma relative to pyroxenes, if diffusion of H2O is faster in olivine. The homogeneous distribution of water in single pyroxene grains and
the relatively constant Dcpx/opx in natural rocks (2 0·5)
indicate that pyroxenes are more likely to preserve the original water content of their mantle source (Peslier et al.,
2002; Peslier & Luhr, 2006; Sundvall, 2010; Xia et al.,
2010), despite apparently similar but slightly lower diffusivities than olivine and only slightly lower than them
(Ingrin & Blanchard, 2006). The Dcpx/opx ratios are in
better agreement in experimental (0·9^3·3) and natural
systems (0·5^6·5) (Table 6), indicating that the partitioning
of water between pyroxenes is less affected by water loss.
Dopx/ol (4·7^6·7) measured in this study, however, is
slightly lower than values reported previously in experimental studies (6·5^25·3). For Dcpx/opx, the experiments in
hydrous basaltic systems (Aubaud et al., 2004; Hauri et al.,
2006; Tenner et al., 2009; Table 6) show lower values
(0·9^1·3) than natural rocks (0·5^6·5) and our study
(2·9^3·3) (Table 6). These differences are possibly due to
higher temperatures, initial higher water concentration
and different bulk chemistry of the hydrous basaltic experiments relative to ours. It should be noted that Al was
demonstrated to have a considerable impact on the
solubility, and therefore the partitioning of water between
NAMs (i.e. Rauch & Keppler, 2002; Stalder, 2004; Hauri
et al., 2006; Tenner et al., 2009; O’Leary et al., 2010). Thus,
it is also important to evaluate the possible impact of differences in Al content of NAMs with respect to other experimental studies. Orthopyroxene and clinopyroxene have
2·03^9·33 and 5·14^14·50 wt % Al2O3 in the studies by
Aubaud et al. (2004), Hauri et al. (2006), Tenner et al. (2009)
and O’Leary et al. (2010) (Table 5). In comparison, orthopyroxene and clinopyroxene in our study have 1·27^2·78
and 1·97^3·67 wt % Al2O3 (Table 5) and are buffered by a
full mantle assemblage consisting of olivine, orthopyroxene, clinopyroxene and garnet. This means that both the
orthopyroxene and the clinopyroxene of our experiments
contain less Al2O3 than the pyroxenes of the aforementioned studies. This lower water content may apply to
orthopyroxene only, as our clinopyroxene has water concentrations comparable with those of these other experimental studies. This indicates that Al2O3 may have an
effect on partitioning of water between orthopyroxene and
clinopyroxene, and orthopyroxene may have more water
if Al2O3 concentration is higher.
The effect of Al2O3 on the substitution of water into olivine could not be addressed in detail with the present dataset, as the EDS measurements for Al2O3 in olivine are not
accurate. The position of the most intense [triv] absoption
bands in olivine is at 3354 cm1 (Table 3). Berry et al.
(2007) showed that the most intense absorption band in
Al2O3-doped olivine occurs at 3345 cm1, which is
beyond any analytical error ( 3 cm1) at lower wavenumber. Our observed [triv] absorption is closer to that of
Fe3þ in olivine at 3350 cm1 (Berry et al., 2007). In contrast
to Hauri et al. (2006), this would seem to exclude Al2O3 as
an important influence on the solubility of water in olivine.
The water content in NAMs coexisting with pargasite,
phlogopite or both was determined in configuration 4 experiments (Table 2). The water contents of pargasite and
phlogopite were assumed to be 1·5 and 4 wt % H2O respectively from stoichiometry, giving Dparg/ol ¼ 3·4 102
and Dparg/opx ¼ 50^52 at 2·5 GPa, 10008C (P12 and P9).
The effect of melting, initial water
concentration and chemistry on water
solubility in peridotitic NAMs
Our experiments at 2·5 GPa at the vapour-saturated solidus of lherzolite show that the solubility of water in olivine
does not decrease at 2·5 GPa when melting begins
(Table 2). Pargasite, however, disappears at or very close
to the solidus (Green, 1973; Niida & Green, 1999; Green
et al., 2010). Similarly, at 4 GPa the water content in olivine
and orthopyroxene does not decrease beyond analytical
uncertainty in two experiments with HZ1 þ5 wt % ‘dry’
phlogopite composition in configuration 4 (Table 2). The
olivine and orthopyroxene from the supra-solidus
2087
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 10
OCTOBER 2012
5.0
4.5
4.0
Bell & Rossman (1992)
Peslier et al. (2002),
3.5
Peslier &Luhr (2006)
Grant et al. (2007a)
3.0
Li et al. (2008)
DHcpx2O/ opx
2.5
Yang et al. (2008)
Falus et al. (2008)
2.0
Aubaud et al. (2004)
1.5
Hauri et al. (2006)
1.0
Tenner et al. (2009)
This study (4 GPa)
0.5
This study (2.5 GPa)
0
0
20
40
D
opx / ol
H 2O
60
80
100
Fig. 9. Partition coefficients of Dcpx/opx versus Dopx/ol for natural upper mantle xenoliths and experimental samples. The bars do not indicate the
error, but the range of partition coefficients between coexisting mineral pairs for a given sample suite. The partititon coefficient values from
this study are based on P8^P12 for Dopx/ol and P6 and P7 for Dcpx/opx inTable 2. Filled black diamond and square are for 4 GPa and 2·5 GPa respectively. The bars for our partition coefficients show the uncertainty arising from the analytical uncertainties in determining the water
concentration.
experiment (62 and 311ppm respectively, P8, 11508C,
Table 2) have, within error, the same water concentration
as their subsolidus counterparts (71 and 317 ppm, P11,
11008C, Table 2). In this latter case phlogopite disappears
at or very close to the solidus (Green et al., 2010). It is
emphasized that the initial appearance of melt does not
‘dry out’ the residual NAMs; this requires high degrees of
melting, as suggested by Hirth & Kohlstedt (1996). The
water concentration in NAMs changes continuously with
increasing temperature across the solidus, decreasing as
melt fraction increases and maintaining the appropriate
partitioning coefficients (possibly changing with temperature as previously noted) between NAMs and the coexisting melt.
The water solubility in olivine from the configuration 1
experiments below the solidus at 4 GPa is 188 ppm when
the initial water concentration is 1·45 wt %, but only
110 ppm with 0·145 wt % (O80 and O92, respectively,
Table 2). The implied lower activity of water in O92 is
attributed to greater concentrations of solutes such as alkalis and contaminating volatile species (such as nitrogen
or carbon dioxide) decreasing the activity of water in the
2088
KOVA¤CS et al.
WATER IN NOMINALLY ANHYDROUS MINERALS
vapour. With lower added water in the experiments,
the relative proportion of entrapped air and carbon
dioxide entering the fluid is enhanced. In contrast, in
configuration 4 experiments (P8, P11, P12, Table 2), where
there is an additional 5 wt % of ‘dry’ phlogopite component in the lherzolite mix, the initial water concentration
(1·45 wt % or 0·145 wt %) does not appear to have
an effect on the water solubility in NAMs within experimental error.
The water content in olivine and orthopyroxene in the
nominally ‘dry’ run at 2·5 GPa (32 and 224 ppm respectively, P30, HZ1) suggests that the water activity is slightly
lower than in vapour-saturated and water-absent experiments at the same pressure (44 and 291^298 ppm, P9 and
P10 respectively, HZ1 þ5 wt % ‘dry’ phlogopite) (Table 2)
but is high enough to stabilize pargasite (the most potassic
observed in the experimental series). The 4 GPa ‘dry’ experiment (P32) in the HZ1 þ5 wt % ‘dry’ phlogopite
composition (Table 2) at 11758C is above the
vapour-saturated solidus for this K-enriched composition
(11008C5Tsolidus511508C). The absence of melting in this
experiment indicates that, at 4 GPa also, the nominally
‘dry’ experiments may produce a lower activity of H2O
than the ‘vapour-saturated’ conditions with 1·45 wt %
H2O present (P11, Table 2). However, the water contents
in olivine and orthopyroxene (57 and 264 ppm respectively), are within analytical uncertainty indistinguishable
from those of its vapour-saturated counterpart (71 and
317 ppm, P11, HZ1 þ5 wt % ‘dry’ phlogopite; Tables 2
and 6).
This ‘dry’experiment (P32 at 4 GPa; phlogopite present)
and those with 0·145 wt % H2O, namely P10 at 2·5 GPa
(pargasite and phlogopite present) and P12 at 4 GPa
(phlogopite present) (Tables 2 and 6), may be referred to
as ‘vapour-absent’. These three vapour-absent experiments
have water concentrations in olivine and orthopyroxene
similar to those of their vapour-saturated counterparts
with 1·45 wt % H2O (P9 and P11 at 2·5 and 4 GPa, Table
2). This implies that in the K-rich experiments the
vapour-absent experiments are in fact close to vapour saturation. All in all, our experimental methods for nominally ‘dry’ experiments did not completely exclude water,
possibly because of hydrogen diffusion through the noble
metal capsule (e.g. Boyd et al., 1964; Chen & Presnall,
1975) or fluid inclusions and water in the olivine and pyroxenes of the mineral layers. It appears that bulk water concentrations lower than 190 ppm cannot be achieved with
the times and techniques used in this study.
It seems that at a given pressure there may be a correlation between the Mg# of olivine and its water content.
Starting composition HZ2 crystallizes higher Mg# olivine than HZ1 (Table 1). At 4 GPa olivine has 62 ppm
(O76) and 188 ppm (O80) water in the HZ2 and HZ1
compositions, respectively (Table 2), although the 508C
difference between the two experiments may play a part.
At 2·5 GPa olivine has 31ppm (O52) and 67 ppm (O85)
water in HZ2 and HZ1 compositions respectively. A similar effect was shown by Withers et al. (2011). At 4 GPa, temperature may also influence olivine H2O storage capacity
as O80 and O81 H2O decreases from 188 to 115 ppm from
1100 to 12008C.
Implications for mantle melting
Our experiments establish that up to 190 ppm H2O can
be accommodated in NAMs in fertile mantle peridotite at
2·5 GPa. Higher bulk H2O contents at this pressure at subsolidus temperatures result in the formation of pargasite,
whose abundance, however, is limited by bulk Na2O contents and other compositional factors as well as temperature and pressure (e.g. Niida & Green, 1999). For fertile
compositions 0·1^0·4 wt % H2O can be accommodated
in pargasite plus NAMs. The solidus of pargasite-bearing
peridotite at vapour-absent conditions is defined by a reaction in which pargasite is replaced by hydrous melt. The
appearance of hydrous melt does not significantly decrease
the water content of NAMs; that is, partial melting does
not ‘dry out’ the residual mantle. The low water contents
in natural lherzolites from the lithosphere (i.e. Peslier,
2010) are consistent with being residual from extraction of
melts with very low water content (such as MORB),
probably indicating a high melt fraction. Alternatively,
low water contents in natural lherzolites could also be
due to subsolidus re-equilibration with vapour with low
water activity (e.g. CO2-rich or CH4-rich). At pressures
greater than 3 GPa pargasite is not stable (Green, 1973;
Niida & Green, 1999; Green et al., 2010) and the storage
capacity of water in fertile subsolidus lherzolite is
190 ppm in NAMs. This water storage capacity could increase significantly if the lherzolite contained enough
K2O to stabilize phlogopite. In the absence of phlogopite
at pressures 43 GPa, then for water contents in excess of
190 ppm, melting occurs at the vapour-saturated solidus
with the melt fraction directly proportional to the excess
water content. Our results suggests that 190 ppm water
is the natural limit for fertile upper mantle at depths
490 km.
The stability of pargasite at very low concentrations has
some important rheological and tectonic implications.
These were explored in the study by Green et al. (2010),
where the lithosphere^asthenosphere rheological boundary was attributed to pargasite breakdown at 3 GPa and
entry of the intraplate geotherm into a region of partial
melting. An observable effect on seismic properties would
also be expected and we note that Thybo (2006) proposed
the presence of a global seismic velocity anomaly at 90^
100 km depth, which, in our interpretation, would coincide
very well with the high-pressure instability of pargasite.
2089
JOURNAL OF PETROLOGY
VOLUME 53
AC K N O W L E D G E M E N T S
We thank Greg Yaxley for fruitful discussions. The authors
would like to acknowledge Frank Brink at the EMU unit
for his assistance with SEM analysis. We thank Anne
Peslier, Trevis Tenner and an anonymous reviewer for comments and M. Wilson for editorial handling of the paper.
F U N DI NG
This research was supported by Australian Research
Council grants to D.H.G. and to G. M. Yaxley and D. H.
Green. I.K. was supported by an A. E. Ringwood
Memorial Scholarship, an Australian International
Postgraduate Research Scholarship and a Marie Curie
International Reintegration Grant (NAMS-230937). A.R.
was supported by an ANU PhD Scholarship.
S U P P L E M E N TA RY DATA
Supplementary data for this paper are available at Journal
of Petrology online.
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