CO2-induced small water solubility in olivine and implications for properties of the shallow mantle

Earth and Planetary Science Letters 403 (2014) 37–47

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CO2 -induced small water solubility in olivine and implications for properties of the shallow mantle Xiaozhi Yang a,b,∗ , Dingding Liu a , Qunke Xia c a

State Key Laboratory for Mineral Deposits Research, School of Earth Sciences and Engineering, Nanjing University, Nanjing 210046, PR China Bayerisches Geoinstitut, Universität Bayreuth, D-95440 Bayreuth, Germany c CAS Key Laboratory of Crust–Mantle Materials and Environments, School of Earth and Space Sciences, University of Science and Technology of China, Hefei 230026, PR China b

a r t i c l e

i n f o

Article history: Received 1 April 2014 Received in revised form 12 June 2014 Accepted 17 June 2014 Available online xxxx Editor: J. Brodholt Keywords: water solubility olivine H2 O–CO2 experimental studies mantle properties

a b s t r a c t H2 O and CO2 are important components of fluids in the mantle at ∼30–150 km depth, and may affect strongly water dissolution in nominally anhydrous olivine; however, available experimental hydrogenation of olivine has been nearly exclusively carried out in coexistence with H2 O (CO2 -free). In this study, the effect of CO2 on water solubility in olivine has been investigated by H-annealing natural olivine under peridotite- and fluid-saturated conditions. Experiments were conducted at 1.5–5 GPa and 1100–1300 ◦ C, with oxygen fugacity controlled by Ni–NiO and with either H2 O or H2 O–CO2 as buffering fluid. The olivine shows no change in composition during the experiments. The infrared spectra of the hydrated olivine are characterized by prominent OH bands from ∼3650 to 3000 cm−1 in all the runs, at both high frequency (>3450 cm−1 ) and low frequency (<3450 cm−1 ), and the H2 O solubility is ∼120–370 ppm for the olivine in coexisting with H2 O, and ∼65–180 ppm for the olivine in coexisting with H2 O–CO2 . When CO2 is present in the buffering fluid, the H2 O solubility of olivine is reduced by a factor of ∼2, due to effect on the partitioning of water between minerals and coexisting fluid, and the measured H2 O solubility shows independence on fluid composition (the molar ratio of CO2 to CO2 + H2 O at ∼0.2–0.5) given pressure, temperature and oxygen fugacity. Olivine equilibrated in the shallow mantle is probably dominated by OH groups in the wavenumber ∼3650–3000 cm−1 , and the intensity of OH bands at low frequency may be higher than or comparable to those at higher frequencies. The storage capacity of water in the shallow mantle in previous estimates may have been overestimated by a factor of at least ∼4 if the observed effect of CO2 on water solubility is correct. Our results have profound influence on understanding partial melting, electrical conductivity anomalies and metasomatism in the shallow mantle. © 2014 Elsevier B.V. All rights reserved.

1. Introduction One of the most important approaches established in mineral physics in the past decades is the recognition that nominally anhydrous minerals, such as olivine and pyroxenes, commonly contain trace amounts of water as H-related point defects in their crystal structure, being main reservoir for water storage in the mantle, and that such water, even on ppm levels, affects strongly some physicochemical properties of the host phases and the mantle, including partial melting and electrical conductivity (e.g. Karato, 1990; Bell and Rossman, 1992; Ingrin and Skogby, 2000; Hirschmann, 2006; Keppler and Smyth, 2006;

*

Corresponding author. E-mail address: [email protected] (X. Yang).

http://dx.doi.org/10.1016/j.epsl.2014.06.025 0012-821X/© 2014 Elsevier B.V. All rights reserved.

Green et al., 2010). Olivine is the most abundant mineral in the upper mantle, and the hydrogenation of olivine, in particular the concentration levels, has received increasing attention as documented by extensive studies on natural and synthetic samples. Because of the much faster H diffusivity in olivine relative to other minerals such as pyroxenes and thus the more significant diffusion loss/gain of water during exhumation (Ingrin and Skogby, 2000; Demouchy and Mackwell, 2006), which in some cases have led to zoned diffusion profiles over grain scales (e.g. Demouchy et al., 2006; Peslier and Luhr, 2006; Peslier et al., 2008), investigation of natural olivines, as captured by volcano-related eruptions, may only provide limited clue about the “actual water” in the mantle. Accordingly, experimental studies of water solubility in olivine at mantle conditions are of crucial importance as they provide important constraints on the storage capacity of water, the partitioning

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X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

of water between coexisting assemblages and partial melting in the mantle, as well as the cycling and exchange of water between Earth’s interior and exterior and other relevant properties of the Earth (e.g. Hirschmann, 2006; Keppler and Smyth, 2006). Experiments of olivine hydrogenation have so far been conducted nearly exclusively at H2 O-saturated conditions, by equilibrating/growing crystals in H2 O or in H2 O-bearing melts (e.g. Bai and Kohlstedt 1992, 1993; Kohlstedt et al., 1996; Matveev et al., 2001; Lemaire et al., 2004; Zhao et al., 2004; Berry et al., 2005, 2007; Mosenfelder et al., 2006; Smyth et al., 2006; Grant et al., 2007; Bali et al., 2008; Withers and Hirschmann, 2008; Ardia et al., 2012; Férot and Bolfan-Casanova, 2012; Kovács et al., 2012; Sokol et al., 2013a, 2013b). These studies have shown that the amount of water dissolved in olivine depends markedly on composition, pressure, temperature, oxygen fugacity and environmental chemistry, e.g. water activity and major assemblages, and the yielded data have been used to explain many properties of the upper mantle. However, fluids in the upper mantle are much more complicated, and they are dominated by H2 O–CO2 in shallow depths and by H2 O–CH4 at greater depths (Wood et al., 1990; Frost and McCammon, 2008; Manning et al., 2013). Thermodynamically, the equation of state (EOS) of H2 O–CO2 or H2 O–CH4 differs from that of H2 O (e.g. Frost and Wood, 1997; Duan and Zhang, 2006; Manning et al., 2013). Consequently, the water activity of fluids in the upper mantle is different from that of H2 O in experimental studies, and this would affect water solubility in olivine. Hence, experimental hydrogenation of olivine under conditions buffered by C–O–H fluids is necessary. In this report, we have for the first time systematically studied the water solubility of olivine coexisting with peridotite and H2 O–CO2 , by conducting experiments at 1.5–5 GPa and 1100–1300 ◦ C, with oxygen fugacity buffered by Ni–NiO and under various fluid compositions concerning the mixing of H2 O and CO2 . For comparison, experiments on water solubility of olivine coexisting with H2 O under otherwise identical conditions have also been undertaken. The purpose of this study was to determine the effect of CO2 on H2 O solubility in olivine and to provide an updated database for water storage in the mantle at ∼30–150 km depths. We show that, when CO2 is present, the water solubility is profoundly reduced. The results are potentially of fundamental implications for a better knowledge of the water storage and partial melting in the shallow mantle, as well as its other properties. 2. The shallow mantle and C–O–H fluids The shallow mantle throughout this paper denotes the mantle portion at depths of ∼30–150 km, and in some continental regions, it coincides roughly with the so-called sub-continental lithospheric mantle. The shallow mantle is dominated by peridotites, and olivine constitutes ∼60% of lherzolite to 90% of dunite peridotites. The shallow mantle is relatively oxidized, with oxygen fugacity mostly within ±2 log units relative to the fayalite– magnetite–quartz (FMQ) buffer, although more reduced/oxidized zones can be present regionally (e.g. Wood et al., 1990; Canil et al., 1994; Bézos and Humler, 2005; Frost and McCammon, 2008; Stagno et al., 2013); in contrast, the deep upper mantle is more reduced, with oxygen fugacity reaching FMQ-4 at ∼200 km (e.g. McCammon, 2005; Rohrbach et al., 2007; Stagno et al., 2013). Although the composition varies substantially between different tectonic settings, C–O–H fluids play significant roles in most mantle geologic processes, such as melting, metasomatism, magma degassing and destruction of continents (e.g. Wood et al., 1990; Lee et al., 2011). The speciation of C–O–H fluids in the mantle is closely related to oxygen fugacity: under oxidizing conditions (e.g. FMQ ± 2), the fluids contain abundant CO2 , while under reducing

Fig. 1. The speciation of C–O–H fluid as a function of pressure in the upper mantle along an adiabat with a potential temperature of 1200 ◦ C, on the basis of potential oxidation state of the upper mantle determined by peridotite assemblages and of prediction that the upper mantle becomes increasing reduced with depth (after Frost and McCammon, 2008). Oxygen fugacity decreases with increasing pressure. At ∼3 GPa, graphite precipitates, and at oxygen fugacities below the CCO buffer (C + O2 = CO2 ), indicated by the vertical line, graphite is unstable and mixtures of CO2 and H2 O are always stable. At higher pressures, fluid speciation is controlled by the CCO buffer.

conditions (e.g. FMQ-4), the fluids are CH4 -rich (Wood et al., 1990; Frost and McCammon, 2008). Fig. 1 shows a calculated speciation model for C–O–H fluids along an upper mantle adiabat (Frost and McCammon, 2008). The most obvious observation is that, above ∼5 GPa, CH4 becomes gradually dominated upon the expense of CO2 . Note, however, that the calculations in Fig. 1 rely on extrapolation of thermodynamic data, so that absolute proportions of fluid species are uncertain, although the general trends are probably reliable. According to Fig. 1, fluids in the shallow mantle are mainly H2 O and CO2 . The presence of quantitatively significant CO2 at such depths has been well documented by the recognition of (1) the global occurrence of CO2 -rich fluid inclusions in the minerals of peridotite xenoliths from a variety of tectonic settings (e.g. Andersen and Neumann, 2001), (2) a distinctive C-, O- and Sr-isotopic signature of mantle origin for some carbonates found in carbonatites and kimberlites (e.g. Deines, 2002), and (3) the wide appearance of CO2 in volcanic/fumarolic gases associated with basaltic eruptions and also found as inclusions within phenocrysts/megacrysts (e.g. Luth, 2003). Fluid inclusions in minerals of mantle xenoliths/xenocrysts represent accidentally trapped samples of fluids present in the shallow mantle, and are thus of unique importance for understanding mantle fluid species. Fluid inclusions in mantle samples are, however, easily subjected to leakage or other secondary processes (e.g. Andersen and Neumann, 2001), so that reconstruction of their initial composition is usually difficult. Analyses of mid-ocean-ridge basalts (MORB) and abyssal peridotites have shown that, by considering the effect of partial melting, crystal fractionation and degassing, typical sub-oceanic mantle contains ∼100 ± 50 ppm H2 O (e.g. 142 ± 85 ppm H2 O, Saal et al., 2002; 116 ± 58 ppm H2 O, Salters and Stracke, 2004; 110 ± 50 ppm H2 O, Workman and Hart, 2005) and ∼60 ± 30 ppm CO2 (e.g. 72 ± 19 ppm CO2 , Saal et al., 2002; 50 ± 12 ppm CO2 , Salters and Stracke, 2004; 36 ± 12 ppm CO2 , Workman and Hart, 2005). For a first-order approximation, these data yield a molar ratio of ∼0.2 for CO2 to H2 O + CO2 , X(CO2 ), in shallow mantle. On the other hand, the source mantle of ocean-island basalts (OIB) or enriched-MORB contains ∼300–900 ppm H2 O and up to

X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

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Table 1 Chemical composition of the starting minerals (wt.%).

Peridotite olivine Peridotite orthopyroxene Peridotite clinopyroxene Peridotite spinel Starting olivine crystal

SiO2

TiO2

Al2 O3

Cr2 O3

FeO

MnO

MgO

CaO

Na2 O

K2 O

NiO

Total

41.39 57.17 54.35 <0.01 41.24

<0.01 0.06 0.20 0.02 <0.01

0.01 2.54 4.20 37.43 0.01

0.01 0.56 1.39 31.17 0.02

8.19 5.47 1.93 12.13 8.34

0.10 0.14 0.08 0.09 0.12

49.83 33.27 15.14 17.75 49.92

0.03 0.45 20.46 0.01 0.03

<0.01 0.10 1.92 <0.01 <0.01

<0.01 <0.01 <0.01 <0.01 <0.01

0.40 0.10 0.06 0.26 0.40

99.95 99.87 99.73 98.87 100.07

FeO: assuming all Fe as FeO. Data are the average of multi-points (5–10) analyses by EMP. The composition is essentially the same between the single crystal olivine and the peridotite olivine, so that equilibrium between the four minerals can be reasonably assumed. Table 2 Summary of experimental conditions and water content in olivine. No.

P (GPa)

T (◦ C)

O-buffer

Fluid-source

Fluid

X(CO2 )

Duration (h)

ppm H2 Oc

ppm H2 Od

PC4a PC4a PC10 PC7 PC3 PC25a PC25a PC22a,b PC22a,b PC23a,b PC23a,b PC13 PC17 PC18 PC19 PC2 PC1a PC1a PC11 PC15 PC16 MA1 MA2

1.5 1.5 1.5 1.5 1.5 1.5 1.5 1.5 1.5 1.5 1.5 2.5 2.5 2.5 2.5 2.5 2.5 2.5 3.5 3.5 3.5 5 5

1100 1100 1100 1100 1100 1100 1100 1100 1100 1100 1100 1100 1100 1250 1250 1300 1300 1300 1100 1100 1100 1100 1100

Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO Ni–NiO

distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled distilled

H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O H2 O

0 0 0 0.22 0.50 0.37 0.37 0.33 0.33 0.45 0.45 0 0.24 0 0.20 0 0.21 0.21 0 0.21 0.23 0 0.22

30 30 30 30 30 30 30 30 30 30 30 30 30 30 30 10 10 10 30 30 30 13.5 13.5

119 127 118 62 65 67 64 63 71 71 68 200 72 316 171 370 180 190 239 114 117 364 146

75 80 75 39 41 42 41 40 45 45 43 127 46 200 108 234 114 120 151 72 74 230 92

water water water water water water water water water water water water water water water water water water water water water water water

+ + + + + + + +

NaHCO3 NaHCO3 NaHCO3 NaHCO3 Ag2 C2 O4 Ag2 C2 O4 Ag2 C2 O4 Ag2 C2 O4

+ NaHCO3 + NaHCO3 + NaHCO3 + NaHCO3 + NaHCO3 + NaHCO3 + NaHCO3

+ + + + + + + +

CO2 CO2 CO2 CO2 CO2 CO2 CO2 CO2

+ CO2 + CO2 + CO2 + CO2 + CO2 + CO2 + CO2

The decomposition of NaHCO3 at high temperature releases H2 O and CO2 , 2NaHCO3 = Na2 CO3 + H2 O + CO2 , and the decomposition of Ag2 C2 O4 upon heating releases CO2 only, Ag2 C2 O4 = 2Ag + 2CO2 . X(CO2 ) is the molar ratio of CO2 to H2 O + CO2 , calculated according to the fluid materials added in the sealed capsule (relative loss of fluid during the runs is not significant: see text). a Two grains from the same run were measured for water content. b Ag2 C2 O4 was used as CO2 source (NaHCO3 was used as CO2 source in the other CO2 -present runs). c Calibrated with the method of Bell et al. (2003). d Calibrated with the method of Withers et al. (2012).

>1000 ppm CO2 (e.g. Javoy and Pineau, 1991; Dixon et al., 2002), but may only constitute a very minor fraction of the shallow mantle. 3. Experimental and analytical methods 3.1. Starting materials The starting materials were a gem-quality single crystal olivine,

∼8 × 12 × 17 mm in size, from Dak Lak (Vietnam), occurring in a lherzolite xenolith within basalt flow, and a fresh lherzolite xenolith from Shuangqing (China), consisting of olivine (∼68%), orthopyroxene (∼23%), clinopyroxene (∼7%), and spinel (∼2%). The composition is homogeneous for each starting mineral, and is essentially the same between the single crystal olivine and peridotite olivine (Table 1), ensuring the approximation of chemical equilibrium between the starting phases. The olivine crystal was cut into ∼1 × 1 × 1.5 mm blocks, and the peridotite was ground into fine powder (mostly <30 μm), which was used to create peridotitesaturated conditions in the shallow mantle and also used as pressure media in the experiments. Starting H2 O content of the olivine crystal was <1 ppm (by weight).

3.2. Experimental studies Experiments were carried out at 1.5–5 GPa and 1100–1300 ◦ C at the Bayerisches Geoinstitut, with a piston cylinder apparatus (1.5–3.5 GPa) and a multi-anvil press (5 GPa). Details of the experiments, including starting materials and working conditions, are given in Table 2. The olivine block was loaded into a Pt capsule (piston cylinder runs: OD 5.0 mm, ID 4.6 mm, length 10.0 mm; multi anvil runs: OD 3.5 mm, ID 3.2 mm, length 5.0 mm), together with peridotite powder, Ni–NiO pairs and fluid-related materials (distilled water/NaHCO3 /Ag2 C2 O4 : Fig. 2), and the capsule was then welded in liquid-N2 cooled environment (weight losses after sealing were <1–2% of the mass of water added). The peridotite powder separated the olivine crystal from the Pt-capsule, as employed in Yang and Keppler (2011), to avoid reaction between them and therefore Fe-loss from the crystal to the capsule, and the Ni–NiO buffers set the oxygen fugacity close to the shallow mantle (Frost and McCammon, 2008). Relative to all the materials sealed in the capsule, the bulk H2 O content, including distilled water added and water released by decomposition of NaHCO3 when loaded, is ∼5–9 wt.% for all the runs, and the bulk CO2 content, excluding the product Na2 CO3 by decomposition of NaHCO3 , is ∼4–6 wt.% for the runs buffered by CO2 + H2 O, with X(CO2 ) of ∼0.2–0.5.

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X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

under a microscope. Compositions of the olivine crystals were measured by electron microprobe (EMP), and textures of the matrix were observed by scanning electron microscopy (SEM). The olivine crystals were double-polished along two perpendicular planes (Fig. 3) for subsequent analyses by Fourier-transform infrared (FTIR) spectroscopy, with thicknesses of ∼0.21–0.78 mm as measured by a digital micrometer, and the matrix of some runs was also checked by X-ray powder diffraction. 3.3. Analytical techniques Fig. 2. Schematic illustration of the experimental capsule (not to scale). The olivine crystal was surrounded by peridotite fine powder (matrix in the recovered capsule in the text), and was not in direct contact with Pt-capsule, O-buffer pairs or CO2 -source materials. Ni–NiO and CO2 -source materials (NaHCO3 /Ag2 C2 O4 ) were always loaded into one end of the capsule.

Piston cylinder experiments were conducted using 0.75 (1.5 GPa) and 0.5 (2.5–3.5 GPa) assemblies, with temperature measured by type-S thermocouples, and multi-anvil experiments were performed using 25/15 assemblies, with temperature measured by type-D thermocouples (for details of assemblies see Keppler and Frost, 2005). At the end of each experiment, the sample was quenched to room temperature within several seconds by switching off the heater, and the pressure was released over 10–40 h to avoid serious cracking of olivine crystals during decompression. The recovered capsules were weighed and pierced, and then heated and reweighed to check for the presence of fluid, and capsules that did not release excess fluid after the runs were discarded. Loss of fluid through Pt capsule in successful experiments was usually negligible, as also documented in some previous reports for runs at relatively oxidizing conditions (e.g. Kohlstedt et al., 1996; Matveev et al., 2001; Sokol et al., 2013a). The capsules were cut open, and olivine crystals were separated from the peridotite matrix. The presence of both Ni and NiO in each capsule was confirmed by examination

Composition of the starting materials and annealed olivine crystals were analyzed by a Shimadzu EPMA-1600 EMP, using 15 kV accelerating voltage, 10 nA beam current, 1 μm beam diameter, counting times of 10 s for the peak and 5 s for the background, and natural and synthetic silicates and oxides as standards. Phases in the matrix and their textures were determined by a JEOL-JSM 6490 SEM. For some samples, diffraction patterns of the matrix were checked with a Rigaku D/MAX-III diffractometer. The general principle for determining water concentration in silicate minerals by FTIR spectroscopy can be described by the Beer–Lambert law: C w = Abstotal / I , where C w is the H2 O content (ppm), Abstotal is the thickness-normalized total integral absorbance (cm−2 ) of OH groups in the mid-IR region, and I is the mineral-specific integral molar absorption coefficient (ppm−1 cm−2 ). Thus, a quantitative evaluation of water content relies critically on the accurate measurement of Abstotal , given a reliable calibration coefficient. For optically anisotropic minerals, intensity of OH absorption bands depends on the orientation of the IR active dipole relative to the incident radiation, and Abstotal is the sum of absorbance along the crystallographic axes (a, b, c), Abstotal = Absa + Absb + Absc . For orthorhombic minerals, Abstotal can be obtained by measuring the absorbance along any three orthogonal directions ( X  , Y  , Z  ) of a crystal, Abstotal = Abs X  + AbsY  + Abs Z  (Libowitzky and Rossman, 1996).

Fig. 3. Examples of experimental products. (a) Polished olivine crystal from run PC2 annealed at 2.5 GPa, 1300 ◦ C and X(CO2 ) = 0 (with distilled water as fluid source). Inset is the SEM image of the matrix from the same run, and minor amounts of melt, occurring as some fluffy materials, is observed at grain boundaries. (b) Polished olivine crystal from run PC7 annealed at 1.5 GPa, 1100 ◦ C and X(CO2 ) = 0.22 (with distilled water + NaHCO3 as fluid source). (c) Decomposed product of NaHCO3 , amorphous Na2 CO3 , from run PC19 annealed at 2.5 GPa, 1250 ◦ C and X(CO2 ) = 0.20 (with distilled water + NaHCO3 as fluid source). (d) SEM image of the matrix from run PC10 annealed at 1.5 GPa, 1100 ◦ C and X(CO2 ) = 0 (with distilled water as fluid source). (e) SEM image of the matrix from run PC22 annealed at 1.5 GPa, 1100 ◦ C and X(CO2 ) = 0.33 (with distilled water + Ag2 C2 O4 as fluid source). (f) SEM image of the matrix from run PC19 annealed at 2.5 GPa, 1250 ◦ C and X(CO2 ) = 0.20 (with distilled water + NaHCO3 as fluid source). The olivine crystals in (a) and (b) were double-polished along two perpendicular planes, and X  , Y  and Z  denote three mutually perpendicular directions along which the polished FTIR spectra were measured (the crystals were slightly tilted for taking the pictures). Some dark-grey materials at grain boundaries or within crystal voids in the SEM images in (d), (e), (f) and inset of (a) are due to epoxy during embedding. The amorphous Na2 CO3 in (c) was separated from the decomposed product of NaHCO3 , located at one end of the capsule (see Fig. 2). oli, olivine; opx, orthopyroxene; cpx, clinopyroxene; sp, spinel.

X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

This has been shown to work even for monoclinic minerals such as feldspars (Johnson and Rossman, 2003). Therefore, Abstotal of OH bands in olivine can be determined from the polarized spectra along any three perpendicular directions of a crystal. This method yields the same C w , given the calibration coefficient, as that by measuring the spectra along the a, b and c axes of oriented crystals, and the only disadvantage is that structural information of OH sites is not available (Rossman, 2014, personal communication). Infrared spectra were acquired using a Bruker IFS 120 spectrometer coupled with a Bruker IR microscope. 100/200 scans were accumulated for each spectrum with 4 cm−1 resolution using a tungsten light source, a Si-coated CaF2 beam-splitter and a narrowband MCT detector. The spot size on the sample was 60 μm in diameter. Polarized radiation was produced by a wire-strip polarizer on a KRS-5 substrate, and polarized spectra were recorded with the electric field vector (E) parallel to three mutually perpendicular directions (Fig. 3). Optically clean and inclusion- and crackfree areas were chosen for the analyses. During the measurements, the optics of the microscope were continuously purged with H2 Oand CO2 -free, purified air. For some representative infrared spectra, several reasonable baseline corrections were performed for background subtraction, and the uncertainty in the integrated total absorbances is usually <10%. A few calibration methods are available for estimating water content in olivine by integrating the infrared spectra over the wavenumber range 3700–3000 cm−1 (see Supplementary material). Two general calibration factors (for multiplying the total absorbance), 0.188 ppm H2 O cm2 by Bell et al. (2003) and 0.119 ppm H2 O cm2 by Withers et al. (2012), were adopted, although it should be mentioned that the application of these coefficients to samples with significant low-frequency OH bands may overestimate the water content because the molar absorptivity of OH bands at low frequencies could be higher (Bell et al., 2003; Kovács et al., 2010), and the calculated results show noticeable difference, implying that the quantification of water in olivine requires further effort. For a safe comparison with many early reported data, the discussions below are based mainly on the method of Bell et al. (2003). 4. Results 4.1. Textures and composition of annealed samples In the recovered peridotite matrix of all experiments, significant grain growth took place, resulting in grains with diameters up to ∼300 μm (Fig. 3). This has also been observed by Mosenfelder et al. (2006) for experiments using finely powdered olivine as starting material. In the runs NaHCO3 was loaded, amorphous carbonatite melt, dominated by Na2 CO3 with some dissolved silicates, was found at one end of the capsules (Fig. 3(c)) and at grain boundaries of the matrix (Fig. 3(f)). In the runs buffered by H2 O, no melt was observed in the sample charges (Fig. 3(d)), except the one at 2.5 GPa and 1300 ◦ C (partial melting occurred at this temperature: inset in Fig. 3(a)). In the runs Ag2 C2 O4 was loaded, no melt was observed in the charges (Fig. 3(e)), and this has also been confirmed by X-ray analyses of the matrix powder, which show the same diffraction patterns as those of mantle olivine, orthopyroxene and clinopyroxene and demonstrate no signals of carbonates (Fig. 4). In the recovered capsules, olivine, orthopyroxene, clinopyroxene and spinel added as fine powder surrounding the olivine crystals were identified (Fig. 3). Resembling previous experimental hydrogenation of natural olivine (Kohlstedt et al., 1996; Grant et al., 2007; Yang and Keppler, 2011), the rectangular shape of the olivine crystals is usually well-preserved, although most of them were fractured during the experiments and decompression. Different from some early reports (Kohlstedt et al., 1996; Matveev et al., 2001; Grant et al., 2007),

41

Fig. 4. Representative X-ray diffraction patterns of the recovered peridotite fine powder. Samples are taken from the two experiments with Ag2 C2 O4 as CO2 source, labeled as PC22 and PC23 in Table 2 (Ni, NiO and Ag were separated from the peridotite powder). Also shown are the diffraction patterns of some natural carbonates, including magnesite (MgCO3 ), dolomite (CaMg(CO3 )2 ) and calcite (CaCO3 ), and mantle olivine (oli), orthopyroxene (opx) and clinopyroxene (cpx), with data from RRUFF (http://rruff.info). Dashed lines show some typical diffraction patterns of the carbonates. The diffraction patterns of the two recovered samples are essentially the same, and agree well with those of mantle olivine, orthopyroxene and clinopyroxene.

however, overgrowths around or reaction zones in the olivine crystals, by exchange with MgO in these studies (either decomposed by brucite or added directly), are not observed, because of the protection of peridotite powder and the absence of MgO in the capsules. The annealed crystals show no compositional zonation in each sample, and the composition is identical to that of the starting material (Table 3), implying that the crystals did not react with H2 O/CO2 to form hydrous phases, e.g. amphibole, and/or carbonates, e.g. magnesite. 4.2. FTIR measurements of olivine crystals Polarized spectra of the hydrated olivines are shown in Fig. 5. In the mid-IR range 3650–3000 cm−1 , the spectra show OH-related absorption bands at ∼3609, 3595, 3579, 3572, 3567, 3546, 3525, 3504, 3478, 3452, 3404, 3389, 3355, 3327, 3220 and 3175 cm−1 , which are typical for many mantle olivines (Miller et al., 1987; Matsyuk and Langer, 2004; Demouchy et al., 2006; Peslier and Luhr, 2006; Li et al., 2008; Peslier et al., 2008, 2010) and have been observed in many experimentally hydrated Fe-bearing olivines (Bai and Kohlstedt 1992, 1993; Zhao et al., 2004; Demouchy and Mackwell, 2006; Grant et al., 2007; Yang and Keppler, 2011; Kovács et al., 2012; Yang, 2012a). The difference in the relative intensity of individual OH bands between these samples results from analyses along different directions, especially for the olivine hydrated at 2.5 GPa and 1250 ◦ C which show, in the range 3650–3550 cm−1 , a main band at ∼3695 cm−1 plus a shoulder at 3572 cm−1 . Bai and Kohlstedt (1992, 1993) divided the OH bands of olivine into group I bands at high frequency (3650–3450 cm−1 ) and group II bands at low frequency (3450–3000 cm−1 ), which is still widely used in many available reports in spite of the complexities involved with the mechanism of H-incorporation (Supplementary

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Table 3 Chemical composition of the recovered olivine crystals (wt.%).

PC4 PC10 PC7 PC3 PC25 PC22 PC23 PC13 PC17 PC18 PC19 PC2 PC1 PC11 PC15 PC16 MA1 MA2

SiO2

TiO2

Al2 O3

Cr2 O3

FeO

MnO

MgO

CaO

NiO

Total

41.53 41.41 41.30 41.64 41.33 41.66 41.56 40.86 41.05 41.32 40.91 41.10 41.42 41.39 40.62 41.16 41.38 41.23

<0.01 0.01 <0.01 0.02 0.01 <0.01 0.01 0.02 0.03 <0.01 <0.01 <0.01 0.05 <0.01 <0.01 0.01 0.02 <0.01

0.01 0.01 0.01 0.02 0.02 0.02 0.01 <0.01 0.02 0.02 0.02 0.01 0.04 0.01 0.01 0.01 0.01 0.02

0.02 <0.01 0.05 0.02 0.03 0.01 0.02 0.01 <0.01 0.01 0.02 0.07 0.04 <0.01 0.02 <0.01 <0.01 <0.01

8.28 8.22 8.29 8.54 8.26 8.15 8.19 8.25 8.40 8.12 8.21 8.48 8.26 8.34 8.17 8.21 8.19 8.61

0.11 0.08 0.08 0.10 0.11 0.10 0.10 0.06 0.07 0.07 0.07 0.12 0.08 0.09 0.13 0.15 0.11 0.10

49.53 49.77 49.72 49.34 49.40 48.58 49.67 49.47 50.16 50.04 49.91 49.20 49.33 49.43 50.31 49.68 50.19 49.78

0.04 0.02 0.03 0.05 0.04 0.05 0.06 0.04 0.04 0.02 0.01 0.05 0.06 0.05 0.04 0.04 0.02 0.03

0.32 0.38 0.38 0.40 0.39 0.37 0.37 0.35 0.37 0.42 0.33 0.35 0.42 0.39 0.36 0.39 0.32 0.43

99.84 99.90 99.87 100.12 99.58 98.94 99.99 99.06 100.14 100.01 99.48 99.37 99.68 99.71 99.65 99.64 100.22 100.20

FeO: assuming all Fe as FeO. Data are the average of multi-points (3–8) analyses by EMP. The contents of Na2 O or K2 O are usually less than 0.01% and are not shown here.

Fig. 5. Polarized FTIR spectra of the recovered olivine crystals as a function of (a) pressure at 1100 ◦ C, (b) temperature at 2.5 GPa, and (c) X(CO2 ) at 1.5 GPa and 1100 ◦ C. In (a) and (b): top panel, with H2 O as the coexisting fluid; bottom panel, with H2 O–CO2 as the coexisting fluid, X(CO2 ) = ∼0.2. The spectra were normalized to 1 cm thickness and vertically offset. Note that a shoulder band at ∼3572 cm−1 is apparent for all the spectra measured at 2.5 GPa and 1250 ◦ C in (b). Different colors (black, red and blue) of the spectra of each sample denote different directions along three mutually perpendicular directions ( X  , Y  and Z  ), which may not be the same between different runs (see text for more details). Dashed lines in each panel mark the positions of Ti-related OH bands at ∼3572 and 3525 cm−1 as proposed by Berry et al. (2005). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

material). In all our experiments, both group I and group II bands are prominent. Profile analyses along core-to-rim paths show no zoned OH distribution (Fig. 6), indicating the attainment of equilibrium H-incorporation. The absence of OH bands representing other minerals such as carbonates (characterized by broad band at ∼3610 cm−1 : Huang and Perr, 1960) or amphibole (characterized by sharp bands above 3620 cm−1 : Skogby and Rossman, 1991), Figs. 5 and 6, suggests that, during the experiments, reactions between fluids and olivine in the charges are unlikely, because FTIR spectroscopy is very sensitive to their products (even appearing as lamellae on nanometer scales). The calculated water contents are given in Table 2. These values represent water solubility under the experimental conditions, as equilibrium hydrogenation was acquired. The solubility generally increases with pressure or temperature, with the system buffered by either H2 O or H2 O–CO2 (Fig. 7(a) and (b)). Enhanced dissolution of water in olivine with pressure has been reported at 2–12 GPa and 1100 ◦ C by Kohlstedt et al. (1996) and Mosenfelder et al. (2006), and with temperature at 1000–1300 ◦ C and 0.3 GPa by Zhao et al. (2004) and at 1100–1400 ◦ C and 2.5 GPa by Bali

et al. (2008), for experiments in equilibrium with H2 O. In coexistence with H2 O, the H2 O solubility increases from ∼120 ppm at 1.5 GPa to 365 ppm at 5 GPa at 1100 ◦ C, and from ∼200 ppm at 1100 ◦ C to 370 ppm at 1300 ◦ C at 2.5 GPa; while in coexistence with H2 O–CO2 (X(CO2 ) of ∼0.2), the H2 O solubility increases from ∼65 ppm at 1.5 GPa to 145 ppm at 5 GPa at 1100 ◦ C, and from ∼72 ppm at 1100 ◦ C to 180 ppm at 1300 ◦ C at 2.5 GPa. Surprisingly, in the CO2 -bearing runs, the H2 O solubility is independent of X(CO2 ) at a given pressure and temperature, e.g., ∼65 ppm for X(CO2 ) of ∼0.2–0.5 at 1.5 GPa and 1100 ◦ C (Fig. 7(c)). The independence of H2 O solubility in olivine on X(CO2 ) has been observed for X(CO2 ) of 0.25–0.7 by Matveev et al. (2001) for experiments at 2 GPa and 1300 ◦ C and with Re–ReO2 as O-buffer, although their focus was on the influence of silica activity (aSiO2 ) on H-species and the effect of CO2 on water solubility was neglected. Similarly, a careful examination of the data in Sokol et al. (2013a) suggests that, for their garnet-free samples, H2 O solubility in olivine demonstrates nearly no dependence on fluid composition even in carbonatite melt + H2 O system, for a range of ∼0.3–0.55 in the molar ratio of CO23− to CO23− + H2 O at 6.3 GPa, 1400 ◦ C and

X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

43

Fig. 6. Representative profile FTIR spectra of the recovered olivine crystals annealed at (a) 1.5 GPa, 1100 ◦ C and X(CO2 ) = 0.2 (run PC7), (b) 5 GPa, 1100 ◦ C and X(CO2 ) = 0 (run MA1), (c) 2.5 GPa, 1300 ◦ C and X(CO2 ) = 0.2 (run PC1), and (d) 1.5 GPa, 1100 ◦ C and X(CO2 ) = 0.45 (run PC23). The spectra were measured with a spot size of 60 μm in diameter, and were normalized to 1 cm thickness and vertically offset. Core-to-rim distance for each crystal is illustrated in each panel. Numbers above each spectrum show the integral absorbance (cm−1 , with an uncertainty of <10% due to background subtraction) and the distance from core (μm in parentheses, with an uncertainty of ∼30 μm).

Fig. 7. Water content of the annealed olivine crystals. (a) Water content vs. pressure, (b) water content vs. temperature, (c) water content vs. X(CO2 ), and (d) the ratio of water H O H O–CO2 content in equilibrium with H2 O (C w2 ) to that in equilibrium with H2 O–CO2 (C w2 ). Solid lines in (a) and (b) from least-squares regression of the data measured under different conditions (in coexisting with H2 O and H2 O–CO2 , respectively), and the yielded equations are: in (a), C w = 21.7 + 66.6 × P (GPa) under H2 O-buffered conditions (r 2 = 0.98), and C w = 20.1 + 25.2 × P (GPa) under H2 O–CO2 -buffered conditions (r 2 = 0.93); in (b), C w = −702.9 + 0.82 × T (◦ C) under H2 O-buffered conditions (r 2 = 0.99), and C w = −570.9 + 0.58 × T (◦ C) under H2 O–CO2 -buffered conditions (r 2 = 0.99). Solid line in (c) is plotted at 65 ppm H2 O, around which the measured water contents cluster for X(CO2 ) ranging from ∼0.2 to 0.5. Dashed lines in (a), (b) and (c), labeled as calculated trend, illustrate the trend of the calculated water contents for olivine crystals equilibrating with H2 O–CO2 fluid, starting with the measured water solubility in the H2 O-buffered systems and under the assumption that water solubility is proportional to water activity which was estimated from the EOS for a binary H2 O–CO2 system (see text and Supplementary material). Uncertainty of water content is assumed 10% relative.

Fe–FeO as O buffer. Our data show that, when CO2 is present in the fluid, the water solubility is reduced by a factor of ∼2 (Fig. 7(d)). 5. On OH groups in mantle olivine The FTIR spectra of the olivine hydrated in equilibrium with either H2 O or H2 O–CO2 show both prominent group I and group II

bands at 1.5–5 GPa, and the intensity of group II bands is usually stronger than or comparable to that of group I bands (Fig. 5). Moreover, the OH bands at ∼3572 and 3525 cm−1 , attributed to Ti-related point defects and present in virtually all H-bearing mantle olivines (Berry et al., 2005), are apparent in nearly all the runs, and the spectra reproduce all the OH bands typical for mantle olivines derived from ∼30 to 150 km depth (e.g. Miller

44

X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

et al., 1987; Matsyuk and Langer, 2004; Demouchy et al., 2006; Peslier and Luhr, 2006; Li et al., 2008; Peslier et al., 2008, 2010). Since, as noted above, the experiments were undertaken under conditions corresponding to the shallow mantle, including pressure, temperature, oxygen fugacity and in particular peridotitesaturated environment and fluid speciation, our results indicate that olivine equilibrated in the shallow mantle should have similar patterns of OH groups. However, the OH bands in mantle olivines, captured as megacrysts/xenocrysts or in peridotite xenoliths by volcanic eruptions, are usually dominated by group I bands, and group II bands are only occasionally observed with much lower intensity, although in some cases their intensity is higher than group I bands (e.g. Miller et al., 1987; Matsyuk and Langer, 2004; Demouchy et al., 2006; Peslier and Luhr, 2006; Li et al., 2008; Peslier et al., 2008, 2010). Understanding the origin for the relative paucity of group II bands in mantle olivines is of crucial importance for insights into processes that may have affected the FTIR features during their residence in the shallow mantle or during their transport to Earth’s surface, and may provide critical information about the retention of original water content in their mantle source. It follows that, if our prediction for the presence of both group I and group II bands in olivine occurring at shallow mantle depths is correct (Fig. 5), the patterns of OH groups observed in mantle olivines must have been modified by some mechanism. A favored explanation was that most naturally exhumed olivines were affected by metasomatism of low aSiO2 media, e.g. carbonatite melt rich in MgO (Matveev et al., 2001). That olivines equilibrating with MgO-rich agent are characterized by group I bands only has been reported by many other experimental and theoretical studies (e.g. Kohlstedt et al., 1996; Mosenfelder et al., 2006; Walker et al., 2007; Withers and Hirschmann, 2008; Otsuka and Karato, 2011). However, a careful examination of all these available reports demonstrates that, for Ti-bearing olivines hydrated/crystallized in equilibrium with low aSiO2 fluid/melt, Ti-related OH bands at ∼3572 and 3525 cm−1 , the two most common and intense bands occurring in essentially all mantle samples, are usually, if not always, absent (either both or one of them). Thus, the dominance of the observed “group I” bands in natural olivines cannot be simply attributed to metasomatism of low aSiO2 media. One possible mechanism accounting for the weak intensity to paucity of group II bands in natural olivines is that the initial water information, with OH patterns similar to those in Fig. 5, was reset during sample ascent to the surface, because of the much faster diffusivity of group II bands than group I bands (Padrón-Navarta et al., 2014) and thus more significant diffusion loss of their primary water, e.g. by exchange with the host magma. If this is true, group I bands in natural mantle samples may be more representative of H-related defects in the shallow mantle, and group II bands, if present, may not be used to properly infer mantle processes. Furthermore, caution must be taken when attempts are made to estimate the H2 O content in the mantle by considering natural olivines dominated by group I bands, as well as the physical influence of water on mantle properties such as rheology and stability of continents. 6. H2 O/CO2 and water storage in the shallow mantle In early experimental olivine hydrogenation with H2 O as buffering fluid (starting with natural olivine of similar composition as in this study), the H2 O solubility measured by Kohlstedt et al. (1996) was ∼400 ppm at ∼2.5 GPa and 1100 ◦ C and ∼1500 ppm at ∼5 GPa and 1100 ◦ C (note that H2 O solubility of >2000 ppm was reported by Mosenfelder et al., 2006 and Otsuka and Karato, 2011 at 5 GPa and 1000 ◦ C) by applying the calibration of Bell et al. (2003), much higher than the values determined here, ∼200 ppm

in the former and ∼365 ppm in the latter case (Table 2, Fig. 7(a)). This may be caused by the experiments conducted under different conditions, e.g. non-peridotite environments in previous reports (with olivine (+ orthopyroxene) + MgO in charges) versus peridotite-saturated environments in this work (with olivine + orthopyroxene + clinopyroxene + spinel in charges), because H2 O solubility varies with mineral composition, populations of point defects and water fugacity of the system which are all dependent on the assemblages of coexisting minerals (Hirschmann et al., 2005). With H2 O–CO2 as buffering fluid, the water solubility increases with increasing temperature at 2.5 GPa even when carbonatite melts are present, e.g. some runs where NaHCO3 was added (Fig. 7(b)). This differs profoundly from the trend observed for systems with silicate melts, which show a decrease in water solubility with increasing temperature at given pressure (above ∼1100 ◦ C), e.g. for Fe-bearing olivine at 2.5–9 GPa (Férot and Bolfan-Casanova, 2012), forsterite at 6–12 GPa (Smyth et al., 2006; Bali et al., 2008), and wadsleyite at 15 GPa (Demouchy et al., 2005). In these studies, the drop of water solubility with temperature is attributed to a reduced water activity in the system. The different trend of our data in the carbonate-melt-bearing runs may be caused by: (1) the decrease of water solubility with temperature is more apparent at higher pressures, and/or (2) the effect of hydrous carbonatite melts + CO2 on water solubility in olivine is different from that of hydrous silicate melts. The H2 O solubility of olivine is reduced by a factor of ∼2 in coexistence with H2 O–CO2 compared to H2 O (Table 2, Fig. 7). This resembles the work of Sokol et al. (2013a) where even if their measured H2 O solubility of olivine buffered by hydrous carbonatite melt is about twice lower than that of olivine buffered by hydrous silicate melt as noted above, due to the relatively strong partitioning of water into carbonatite melt. Therefore, one may think that the smaller water solubility in our H2 O–CO2 -bearing experiments is caused by the presence of Na2 CO3 melt. This possibility is, however, not supported by the similar water solubility between the runs with Ag2 C2 O4 as CO2 source material, where carbonatite melt is absent, and the runs with NaHCO3 as CO2 source material, where carbonatite melt is present, at 1.5 GPa and 1100 ◦ C (Fig. 7(c)). As a result, the most likely explanation is that, when CO2 is present, the partitioning of water between olivine and fluid is affected, with a stronger partitioning of water into coexisting H2 O–CO2 than into H2 O. Ultimately, the small H2 O solubility of olivine in coexisting with H2 O–CO2 may be related to the reduced water activity due to the addition of CO2 (e.g. Frost and Wood, 1997; Duan and Zhang, 2006), if we consider that, since similar frequency and shape of OH bands, representing similar OH point-defects, are observed in our samples (Fig. 5), significant change of H substitution mechanism between these runs is not likely, and that, for olivine, the water solubility is proportional to water activity in the system (e.g. Keppler and Bolfan-Casanova, 2006). However, the calculated solubility of olivine coexisting with H2 O–CO2 at various conditions, by applying the EOS for a binary H2 O–CO2 system from Duan and Zhang (2006) (see Supplementary material) and the measured solubility at X(CO2 ) of 0 from this study, is markedly different from those determined (Fig. 7(a), (b) and (c)). This may be related to the highly non-ideal behaviors of H2 O and CO2 and the fact that they are completely miscible and can dissolve some silicates at high pressure and temperature (see review by Manning et al., 2013), so that the water activity in peridotite-buffered experiments differs from that suggested by the EOS for simple H2 O–CO2 systems. In particular for the runs with various X(CO2 ), the nearly constant H2 O solubility, Fig. 7(c), is inconsistent with a gradual decrease of water activity with increasing X(CO2 ) indicated by the EOS (Supplementary material), and it appears that the water activity was

X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

buffered constant, despite the relative amount of CO2 and H2 O, by some unknown mechanism. Alternatively, some other mechanism must have been operative for water dissolution in olivine equilibrating with peridotite and H2 O–CO2 . One of the initial purposes of this study was to quantitively establish an equation modeling the thermodynamics of water dissolution in olivine in coexisting with H2 O–CO2 under shallow mantle conditions, e.g. the so-called solubility law:

 C w = A · f Hn2 O · exp −

P · V



R·T

where A is a temperature-dependent constant, f H2 O is water fugacity, n is an exponent,  V is the volume change of the host structure due to H incorporation, R is the gas constant, P and T are pressure and temperature, respectively. In previous experiments buffered by H2 O, these factors have been modeled for olivine (Kohlstedt et al., 1996; Mosenfelder et al., 2006), garnet (Lu and Keppler, 1997), pyroxenes (Bromiley et al., 2004; Mierdel et al., 2007), forsterite (Bali et al., 2008) and feldspar (Yang, 2012b). Unfortunately, no known EOS can be applied to the peridotitesaturated H2 O–CO2 system, as noted above, and a quantitative modeling is not possible. In any case, however, the small H2 O solubility of olivine in the peridotite-saturated H2 O–CO2 -buffered system, about two times lower than that in the peridotite-saturated H2 O-buffered system (this study) and more than four times lower than that in simple system buffered by olivine (+ orthopyroxene + MgO) + H2 O (e.g. Kohlstedt et al., 1996; Mosenfelder et al., 2006; Otsuka and Karato, 2011), indicates that the water storage capacity of the shallow mantle in previous estimates (e.g. Hirschmann et al., 2005), using the data of Kohlstedt et al. (1996) and Mosenfelder et al. (2006), is probably overestimated by a factor of at least 4, given that, at shallow mantle conditions, the partition coefficient of water between pyroxenes and olivine is ∼10 (Supplementary material), as determined by experimental studies (Koga et al., 2003; Aubaud et al., 2004; Hauri et al., 2006; Kovács et al., 2012) and summarized by Hirschmann et al. (2005). 7. Implications for partial melting in the shallow mantle and other properties At given thermodynamic conditions, H2 O that is in excess of the solubility of a mineral must be present in a fluid, which can be H2 O-rich, silicate-rich or a supercritical phase. The small water solubility of olivine in coexisting with peridotite and H2 O–CO2 , Fig. 7, is important for understanding partial melting in the shallow mantle. Following a typical oceanic geotherm (Turcotte and Schubert, 2002), temperature increases from ∼1100 to 1300 ◦ C over the pressure range from ∼2.5 to 5 GPa (e.g. ∼70 to 150 km depth), and the H2 O solubility of olivine increases from ∼70 to 350 ppm H2 O (as obtained from extrapolation of the data in Fig. 7(a) and (b), which should be even smaller in continental shield regions due to lower temperature, ∼1100 ◦ C according to geotherm). Assuming that the shallow mantle consists of ∼60% olivine and 40% pyroxenes, and that the H2 O partition coefficient is ∼10 between pyroxenes and olivine, as used in the calculations of Hirschmann et al. (2005), Ardia et al. (2012) and Tenner et al. (2012), the estimated bulk H2 O content required to trigger partial melting of peridotites at such depths is ∼320 to 1600 ppm. This is much greater than the ∼100 ± 50 ppm present in the shallow mantle as sampled by MORB (e.g. Saal et al., 2002; Salters and Stracke, 2004; Workman and Hart, 2005), indicating that pervasive hydrous melting at these depths is unlikely. In regions of high H2 O abundance, where ∼300–900 ppm in the source domains of OIB and enriched-MORB may be found (e.g. Javoy and Pineau, 1991; Dixon et al., 2002), local hydrous partial melting may be initiated.

45

This is consistent with the prediction of Ardia et al. (2012) and Férot and Bolfan-Casanova (2012). In a recent publication, Green et al. (2010) proposed a model where hydrous partial melting could occur in shallow mantle containing ∼180 ppm H2 O, according to their measured similar bulk contents of water retained in olivine and pyroxenes coexisting with a presumed fluid-saturated melt at 2.5 GPa, 1025 ◦ C and 4 GPa, 1210 ◦ C. Given the bulk water content and the assumption for mantle modal composition and water partitioning between coexisting minerals, as mentioned before, the H2 O solubility of olivine inferred by Green et al. (2010) is ∼40 ppm at both 2.5 and 4 GPa (for details see also Kovács et al., 2012), which is a prerequisite for their model. The general trend of our data for peridotite-saturated H2 O–CO2 -buffered runs, Fig. 7(b), yields a value of ∼30 ppm at 2.5 GPa and 1025 ◦ C, which agrees well with that of Green et al. (2010) if we consider the uncertainty of data extrapolation and FTIR analyses (note that unpolarized method was used in Green et al. (2010), leading to relatively larger uncertainty of their measured contents). By contrast, the H2 O solubility at 4 GPa and 1210 ◦ C, suggested by our data (Fig. 7(a) and (b)), is greater than that at 3.5 GPa and 1100 ◦ C, ∼115 ppm, which exceeds profoundly the value of ∼40 ppm. Thus, our results do not support the hydrous partial melting model of asthenosphere promoted by Green et al. (2010). Anomalous domains of high electrical conductivity, ∼0.01– 0.1 S/m or higher, and apparent electrical anisotropy, by a factor of ∼2 to >10, have been resolved below ∼50 km depth in the shallow mantle beneath oceanic and continental regions (e.g. Boerner et al., 1999; Evans et al., 2005; Naif et al., 2013). These anomalies were frequently interpreted by water dissolved in olivine due to H-enhanced electrical conductivity and orientation-related electrical anisotropy, following the theoretical approach of Karato (1990). The enhancement of water on olivine electrical conductivity has been proved by experimental studies (e.g. Wang et al., 2006; Yoshino et al., 2009; Poe et al., 2010; Yang, 2012a). In these reports, Wang et al. (2006) and Yoshino et al. (2009) measured the water content of their polycrystalline samples by unpolarized FTIR analyses so that large uncertainty may be involved as noted in our previous work (Yang, 2012a; Yang et al., 2012); in contrast, Poe et al. (2010) and Yang (2012a) determined the water content of their single crystal samples by polarized FTIR analyses so that more accurate data are produced, which suggest that, to account for the observed high conductivity, >600 ppm H2 O is required. This value is far beyond the storage capacity of olivine at equivalent conditions in the shallow mantle based on our data, e.g. <350 ppm H2 O. Hence, it seems unlikely that water in olivine produces the resolved high conductivity in shallow mantle. Moreover, the results of Poe et al. (2010) and Yang (2012a) also show that, with H2 O content <600 ppm, the orientation-related electrical anisotropy of olivine is negligible, which cannot cause the measured electrical anisotropy. Therefore, our data place boundary conditions concerning the H-enhanced electrical conductivity of olivine, and suggest that H-bearing olivine is unable to account for many reported electrical anomalies in the shallow mantle. Possible candidates for the anomalies may include partial melts, e.g. by incipient melting (Sifré et al., 2014), and locally abundant Feand H-rich pyroxenites (Yang and McCammon, 2012). Finally, our results potentially shed new light on the relationship between water in mantle minerals and metasomatism. Metasomatism has been recognized in xenoliths from the oceanic and continental lithosphere and in a range of volcanic settings (e.g. Bailey, 1982; Roden and Murthy, 1985). In the shallow mantle, metasomatism is commonly attributed to infiltration/percolation of H2 O–CO2 -bearing silicate melts or fluids, which become completely miscible at high pressure and temperature (Shen and Keppler, 1997; Bureau and Keppler, 1999; Manning et al., 2013),

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X. Yang et al. / Earth and Planetary Science Letters 403 (2014) 37–47

and carbonatite melts may be effective metasomatic agents only in some areas and limited to relatively high pressures (Frost, 2006; and references therein). In available reports, high H2 O contents of minerals in peridotite xenoliths are nearly exclusively attributed to metasomatism by fluids/silicate melts (Li et al., 2008; Yang et al., 2008; Peslier et al., 2012), and low water contents to dehydration due to metasomatism by oxidized fluids/melts (Peslier et al., 2002) or due to reheating by upwelling (Yang et al., 2008; Xia et al., 2010). Our data show that, if the metasomatic media is H2 O–CO2 -bearing, as commonly observed in the shallow mantle, it may largely extract water once it invades peridotites due to the strong partitioning of water into coexisting H2 O–CO2 , similar to the scenario proposed for the deep upper mantle that flux of carbonatite melts can greatly drive water out of peridotites (Sokol et al., 2013a). For example, given a H2 O content of ∼150 ppm in olivine equilibrated in the mantle at 2.5 GPa and 1100 ◦ C, influx of H2 O–CO2 -bearing fluids/melts would decrease the maximum solubility to ∼72 ppm H2 O (Fig. 7). This provides a novel mechanism to dehydrate the shallow mantle under isothermal–isobaric conditions. Acknowledgements X.Y. thanks Andreas Audétat and Qicheng Fan for supplying the starting olivine crystal from Dak Lak, Vietnam, and the starting peridotite xenolith from Shuangqing, north China, and Geeth Manthilake, Hongzhan Fei, Hans Keppler and Juan Li for technical assistance with multi-anvil experiments, FTIR measurements and SEM imaging. X.Y. communicated with George Rossman, Hans Keppler, Joseph Smyth, Shun Karato and Dan Frost for clarifying some issues. Constructive comments by three anonymous reviewers helped to improve the manuscript. This study was supported by the National Natural Science Foundation of China (41372041), the National Basic Research Program of China (973 Project: 2014CB845904), the Recruitment Program of Global Young Experts (PR China) and the Bayerisches Geoinstitut Visitors Program. Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2014.06.025. References Andersen, T., Neumann, E.R., 2001. Fluid inclusions in mantle xenoliths. Lithos 55, 301–320. Ardia, P., Hirschmann, M.M., Withers, A.C., Tenner, T.J., 2012. H2 O storage capacity of olivine at 5–8 GPa and consequences for dehydration partial melting of the upper mantle. Earth Planet. Sci. Lett. 345–348, 104–116. Aubaud, C., Hauri, E.H., Hirschmann, M.M., 2004. Hydrogen partition coefficients between nominally anhydrous minerals and basaltic melts. Geophys. Res. Lett. 31, L20611. Bai, Q., Kohlstedt, D.L., 1992. Substantial hydrogen solubility in olivine and implications for water storage in the mantle. Nature 357, 672–674. Bai, Q., Kohlstedt, D.L., 1993. Effects of chemical environment on the solubility and incorporation mechanism for hydrogen in olivine. Phys. Chem. Miner. 19, 460–471. Bailey, D.K., 1982. Mantle metasomatism – continuing chemical change within the Earth. Nature 296, 525–530. Bali, E., Bolfan-Casanova, N., Koga, K.T., 2008. Pressure and temperature dependence of H solubility in forsterite: an implication to water activity in the Earth interior. Earth Planet. Sci. Lett. 268, 354–363. Bell, D.R., Rossman, G.R., 1992. Water in Earth’s mantle: the role of nominally anhydrous minerals. Science 255, 1391–1397. Bell, D.R., Rossman, G.R., Maldener, J., Endisch, D., Rauch, F., 2003. Hydroxide in olivine: a quantitative determination of the absolute amount and calibration of the IR spectrum. J. Geophys. Res. 108, B22105. http://dx.doi.org/10.1029/ 2001JB000679. Berry, A.J., Hermann, J., O’Neill, H.S.C., Foran, G.J., 2005. Fingerprinting the water site in mantle olivine. Geology 33, 869–872.

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