IR calibrations for water determination in olivine,r-GeO<Subscript>2</Subscript>, and SiO<Subscript>2</Subscript> polymorphs

Mineral-specific IR absorption coefficients were calculated for natural and synthetic olivine, SiO₂ polymorphs, and GeO₂ with specific isolated OH point defects using quantitative data from independent techniques such as proton–proton scattering, confocal Raman spectroscopy, and secondary ion mass spectrometry. Moreover, we present a routine to detect OH traces in anisotropic minerals using Raman spectroscopy combined with the “Comparator Technique”. In case of olivine and the SiO₂ system, it turns out that the magnitude of ε for one structure is independent of the type of OH point defect and therewith the peak position (quartz ε = 89,000 ± 15,000  \textl \textmol_{\textH2\textO}^-1 \textcm^-2\text{l}\,\text{mol}_{{\text{H}_2}\text{O}}^{-1}\,\text{cm}^{-2}), but it varies as a function of structure (coesite ε = 214,000 ± 14,000  \textl \textmol_{\textH2\textO}^-1 \textcm^-2\text{l}\,\text{mol}_{{\text{H}_2}\text{O}}^{-1}\,\text{cm}^{-2}; stishovite ε = 485,000 ± 109,000  \textl \textmol_{\textH2\textO}^-1 \textcm^-2\text{l}\,\text{mol}_{{\text{H}_2}\text{O}}^{-1}\,\text{cm}^{-2}). Evaluation of data from this study confirms that not using mineral-specific IR calibrations for the OH quantification in nominally anhydrous minerals leads to inaccurate estimations of OH concentrations, which constitute the basis for modeling the Earth’s deep water cycle.     

Infrared spectroscopic characterization of dehydration and accompanying phase transition behaviors in NAT-topology zeolites

Relative humidity ( P_{\textH 2 \textO} P_{{{\text{H}}_{ 2} {\text{O}}}} , partial pressure of water)-dependent dehydration and accompanying phase transitions in NAT-topology zeolites (natrolite, scolecite, and mesolite) were studied under controlled temperature and known P_{\textH 2 \textO} P_{{{\text{H}}_{ 2} {\text{O}}}} conditions by in situ diffuse-reflectance infrared Fourier transform spectroscopy and parallel X-ray powder diffraction. Dehydration was characterized by the disappearance of internal H₂O vibrational modes. The loss of H₂O molecules caused a sequence of structural transitions in which the host framework transformation path was coupled primarily via the thermal motion of guest Na⁺/Ca²⁺ cations and H₂O molecules. The observation of different interactions of H₂O molecules and Na⁺/Ca²⁺ cations with host aluminosilicate frameworks under high- and low- P_{\textH 2 \textO} P_{{{\text{H}}_{ 2} {\text{O}}}} conditions indicated the development of different local strain fields, arising from cation–H₂O interactions in NAT-type channels. These strain fields influence the Si–O/Al–O bond strength and tilting angles within and between tetrahedra as the dehydration temperature is approached. The newly observed infrared bands (at 2,139 cm⁻¹ in natrolite, 2,276 cm⁻¹ in scolecite, and 2,176 and 2,259 cm⁻¹ in mesolite) result from strong cation–H₂O–Al–Si framework interactions in NAT-type channels, and these bands can be used to evaluate the energetic evolution of Na⁺/Ca²⁺ cations before and after phase transitions, especially for scolecite and mesolite. The 2,176 and 2,259 cm⁻¹ absorption bands in mesolite also appear to be related to Na⁺/Ca²⁺ order–disorder that occur when mesolite loses its Ow4 H₂O molecules.     

The H<Subscript>2</Subscript>O solubility of alkali basaltic melts: an experimental study

Experiments were conducted to determine the water solubility of alkali basalts from Etna, Stromboli and Vesuvius volcanoes, Italy. The basaltic melts were equilibrated at 1,200°C with pure water, under oxidized conditions, and at pressures ranging from 163 to 3,842 bars. Our results show that at pressures above 1 kbar, alkali basalts dissolve more water than typical mid-ocean ridge basalts (MORB). Combination of our data with those from previous studies allows the following simple empirical model for the water solubility of basalts of varying alkalinity and fO₂ to be derived: \textH ₂ \textO( \textwt% ) = \text H ₂ \textO_\textMORB ( \textwt% ) + ( 5.84 ×10 ^{- 5} *\textP - 2.29 ×10 ^{- 2} ) ×( \textNa₂ \textO + \textK₂ \textO )( \textwt% ) + 4.67 ×10 ^{- 2} ×\Updelta \textNNO - 2.29 ×10 ^{- 1} {\text{H}}_{ 2} {\text{O}}\left( {{\text{wt}}\% } \right) = {\text{ H}}_{ 2} {\text{O}}_{\text{MORB}} \left( {{\text{wt}}\% } \right) + \left( {5.84 \times 10^{ - 5} *{\text{P}} - 2.29 \times 10^{ - 2} } \right) \times \left( {{\text{Na}}_{2} {\text{O}} + {\text{K}}_{2} {\text{O}}} \right)\left( {{\text{wt}}\% } \right) + 4.67 \times 10^{ - 2} \times \Updelta {\text{NNO}} - 2.29 \times 10^{ - 1} where H₂O_MORB is the water solubility at the calculated P, using the model of Dixon et al. (1995). This equation reproduces the existing database on water solubilities in basaltic melts to within 5%. Interpretation of the speciation data in the context of the glass transition theory shows that water speciation in basalt melts is severely modified during quench. At magmatic temperatures, more than 90% of dissolved water forms hydroxyl groups at all water contents, whilst in natural or synthetic glasses, the amount of molecular water is much larger. A regular solution model with an explicit temperature dependence reproduces well-observed water species. Derivation of the partial molar volume of molecular water using standard thermodynamic considerations yields values close to previous findings if room temperature water species are used. When high temperature species proportions are used, a negative partial molar volume is obtained for molecular water. Calculation of the partial molar volume of total water using H₂O solubility data on basaltic melts at pressures above 1 kbar yields a value of 19 cm³/mol in reasonable agreement with estimates obtained from density measurements.     

The legacy of crystal-plastic deformation in olivine: high-diffusivity pathways during serpentinization

Crystal-plastic olivine deformation to produce subgrain boundaries composed of edge dislocations is an inevitable consequence of asthenospheric mantle flow. Although crystal-plastic deformation and serpentinization are spatio-temporally decoupled, we identified compositional readjustments expressed on the micrometric level as a striped Fe-enriched ( [`(X)]<sub>\textFe</sub> \bar{X}_{\text{Fe}}  = 0.24 ± 0.02 (zones); 0.12 ± 0.02 (bulk)) or Fe-depleted ( [`(X)]_\textFe \bar{X}_{\text{Fe}}  = 0.10 ± 0.01 (zones); 0.13 ± 0.01 (bulk)) zoning in partly serpentinized olivine grains from two upper mantle sections in Norway. Focused ion beam sample preparation combined with transmission electron microscopy (TEM) and aberration-corrected scanning TEM, enabling atomic-level resolved electron energy-loss spectroscopic line profiling, reveals that every zone is immediately associated with a subgrain boundary. We infer that the zonings are a result of the environmental Fe²⁺Mg₋₁ exchange potential during antigorite serpentinization of olivine and the drive toward element exchange equilibrium. This is facilitated by enhanced solid-state diffusion along subgrain boundaries in a system, which otherwise re-equilibrates via dissolution-reprecipitation. Fe enrichment or depletion is controlled by the silica activity imposed on the system by the local olivine/orthopyroxene mass ratio, temperature and the effect of magnetite stability. The Fe-Mg exchange coefficients K_\textD^{\textAtg/\textOl} K_{\text{D}}^{{{\text{Atg}}/{\text{Ol}}}} between both types of zoning and antigorite display coalescence toward exchange equilibrium. With both types of zoning, Mn is enriched and Ni depleted compared with the unaffected bulk composition. Nanometer-sized, heterogeneously distributed antigorite precipitates along olivine subgrain boundaries suggest that water was able to ingress along them. Crystallographic orientation relationships gained via electron backscatter diffraction between olivine grain domains and different serpentine vein generations support the hypothesis that serpentinization was initiated along olivine subgrain boundaries.     

Solubility of alumina in orthopyroxene plus spinel as a geobarometer in complex systems. Applications to spinel-bearing alpine-type peridotites

The addition of Fe and Cr to the simple system MgO-SiO₂-Al₂O₃ markedly affects the activities of phases involved in the equilibrium

H<Subscript>2</Subscript>O diffusion in peralkaline to peraluminous rhyolitic melts

The diffusion of water in a peralkaline and a peraluminous rhyolitic melt was investigated at temperatures of 714–1,493 K and pressures of 100 and 500 MPa. At temperatures below 923 K dehydration experiments were performed on glasses containing about 2 wt% H₂O _t in cold seal pressure vessels. At high temperatures diffusion couples of water-poor (<0.5 wt% H₂O _t ) and water-rich (~2 wt% H₂O _t ) melts were run in an internally heated gas pressure vessel. Argon was the pressure medium in both cases. Concentration profiles of hydrous species (OH groups and H₂O molecules) were measured along the diffusion direction using near-infrared (NIR) microspectroscopy. The bulk water diffusivity () was derived from profiles of total water () using a modified Boltzmann-Matano method as well as using fittings assuming a functional relationship between and Both methods consistently indicate that is proportional to in this range of water contents for both bulk compositions, in agreement with previous work on metaluminous rhyolite. The water diffusivity in the peraluminous melts agrees very well with data for metaluminous rhyolites implying that an excess of Al₂O₃ with respect to alkalis does not affect water diffusion. On the other hand, water diffusion is faster by roughly a factor of two in the peralkaline melt compared to the metaluminous melt. The following expression for the water diffusivity in the peralkaline rhyolite as a function of temperature and pressure was obtained by least-squares fitting:

Systematic study of hydrogen incorporation into Fe-free wadsleyite

The H₂O content of wadsleyite were measured in a wide pressure (13–20 GPa) and temperature range (1,200–1,900°C) using FTIR method. We confirmed significant decrease of the H₂O content of wadsleyite with increasing temperature and reported first systematic data for temperature interval of 1,400–1,900°C. Wadsleyite contains 0.37–0.55 wt% H₂O at 1,600°C, which may be close to its water storage capacity along average mantle geotherm in the transition zone. Accordingly, water storage capacity of the average mantle in the transition zone may be estimated as 0.2–0.3 wt% H₂O. The H₂O contents of wadsleyite at 1,800–1,900°C are 0.22–0.39 wt%, indicating that it can store significant amount of water even under the hot mantle environments. Temperature dependence of the H₂O content of wadsleyite can be described by exponential equation C_{\textH2 \textO} = 6 3 7.0 7 \texte ^{- 0.00 4 8T} , C_{{{\text{H}}_{2} {\text{O}}}} = 6 3 7.0 7 {\text{e}}^{ - 0.00 4 8T} , where T is in °C. This equation is valid for temperature range 1,200–2,100°C with the coefficient of determination R ² = 0.954. Temperature dependence of H₂O partition coefficient between wadsleyite and forsterite (D _wd/fo) is complex. According to our data apparent D_wd/fo decreases with increasing temperature from D _wd/fo = 4–5 at 1,200°C, reaches a minimum of D _wd/fo = 2.0 at 1,400–1,500°C, and then again increases to D _wd/fo = 4–6 at 1,700–1,900°C.     

Water and Iron effect on the P-T-x coordinates of the 410-km discontinuity in the Earth upper mantle

We performed multi-anvil experiments in the system MgO-SiO₂ ± H₂O at 13.0–13.7 GPa and 1,025–1,300°C and in the system MgO-FeO-SiO₂ ± H₂O, under reducing conditions, at 11.0–12.7 GPa and 1,200°C, to depict the effect of H₂O on the P-T-x coordinates of the 410-km discontinuity, i.e. the olivine–wadsleyite phase boundary. The charges were investigated with Electron Microprobe (EMP), Raman Spectroscopy, Fourier Transform Infrared Spectroscopy (FTIR), Secondary Ion Mass Spectrometry (SIMS) and Electron Energy Loss Spectroscopy (EELS). We observe in the MgO-SiO₂-H₂O system at 1,200°C a 0.6 GPa shift of the phase boundary to lower pressure compared to dry conditions, due to the stronger water fractionation into wadsleyite (wad) rather than in olivine (ol). In the MgO-FeO-SiO₂-H₂O system, we reproduced the triple point, i.e. observed coexisting hydrous ol, wad and ringwoodite (ring). SIMS H quantifications provided partitioning coefficients for water: D_\textwad/ol^\textwater D_{\text{wad/ol}}^{\text{water}}  ~ 3.7(5) and D_\textring/ol^\textwater D_{\text{ring/ol}}^{\text{water}}  ~ 1.5(2) and D_{\textwad/ring}^\textwater D_{\text{wad/ring}}^{\text{water}}  ~ 2.5(5). For a bulk composition of x _Fe = 0.1, our data indicate only a slight difference in the width of the loop of the two phase field ol–wad under hydrous conditions compared to dry conditions, i.e. no broadening with respect to composition but a shift to lower pressures. For bulk compositions of x _Fe > 0.2, i.e. in regions where wad–ring and ol–ring coexist, we observe, however, an unexpected broadening of the loops with a shift to higher iron contents. In total, the stability field of hydrous wad expands in both directions, to lower and higher pressures. Fe³⁺ concentrations as determined by EELS are very low and are expected to play no role in the broadening of the loops.     

Experimental and computational study of trace element distribution between orthopyroxene and anhydrous silicate melt: substitution mechanisms and the effect of iron

Although orthopyroxene (Opx) is present during a wide range of magmatic differentiation processes in the terrestrial and lunar mantle, its effect on melt trace element contents is not well quantified. We present results of a combined experimental and computational study of trace element partitioning between Opx and anhydrous silicate melts. Experiments were performed in air at atmospheric pressure and temperatures ranging from 1,326 to 1,420°C in the system CaO–MgO–Al₂O₃–SiO₂ and subsystem CaO–MgO–SiO₂. We provide experimental partition coefficients for a wide range of trace elements (large ion lithophile: Li, Be, B, K, Rb, Sr, Cs, Ba, Th, U; rare earth elements, REE: La, Ce, Nd, Sm, Y, Yb, Lu; high field strength: Zr, Nb, Hf, Ta, Ti; transition metals: Sc, V, Cr, Co) for use in petrogenetic modelling. REE partition coefficients increase from $ D_{\text{La}}^{{\text{Opx}} {\hbox{-}} {\text{melt}}} \sim 0.0005 Although orthopyroxene (Opx) is present during a wide range of magmatic differentiation processes in the terrestrial and lunar mantle, its effect on melt trace element contents is not well quantified. We present results of a combined experimental and computational study of trace element partitioning between Opx and anhydrous silicate melts. Experiments were performed in air at atmospheric pressure and temperatures ranging from 1,326 to 1,420°C in the system CaO–MgO–Al₂O₃–SiO₂ and subsystem CaO–MgO–SiO₂. We provide experimental partition coefficients for a wide range of trace elements (large ion lithophile: Li, Be, B, K, Rb, Sr, Cs, Ba, Th, U; rare earth elements, REE: La, Ce, Nd, Sm, Y, Yb, Lu; high field strength: Zr, Nb, Hf, Ta, Ti; transition metals: Sc, V, Cr, Co) for use in petrogenetic modelling. REE partition coefficients increase from $ D_{\text{La}}^{{\text{Opx}} {\hbox{-}} {\text{melt}}} \sim 0.0005 $ D_{\text{La}}^{{\text{Opx}} {\hbox{-}} {\text{melt}}} \sim 0.0005 to $ D_{\text{Lu}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.109 $ D_{\text{Lu}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.109 , D values for highly charged elements vary from $ D_{\text{Th}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.0026 $ D_{\text{Th}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.0026 through $ D_{\text{Nb}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.0033 $ D_{\text{Nb}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.0033 and $ D_{\text{U}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.0066 $ D_{\text{U}}^{{{\text{Opx}} {\hbox{-}} {\text{melt}}}} \sim 0.0066 to $ D_{\text{Ti}}^{{\text{Opx}} {\hbox{-}} {\text{melt}}} \sim 0.058 $ D_{\text{Ti}}^{{\text{Opx}} {\hbox{-}} {\text{melt}}} \sim 0.058 , and are all virtually independent of temperature. Cr and Co are the only compatible trace elements at the studied conditions. To elucidate charge-balancing mechanisms for incorporation of REE into Opx and to assess the possible influence of Fe on Opx-melt partitioning, we compare our experimental results with computer simulations. In these simulations, we examine major and minor trace element incorporation into the end-members enstatite (Mg₂Si₂O₆) and ferrosilite (Fe₂Si₂O₆). Calculated solution energies show that R²⁺ cations are more soluble in Opx than R³⁺ cations of similar size, consistent with experimental partitioning data. In addition, simulations show charge balancing of R³⁺ cations by coupled substitution with Li⁺ on the M1 site that is energetically favoured over coupled substitution involving Al–Si exchange on the tetrahedrally coordinated site. We derived best-fit values for ideal ionic radii r ₀, maximum partition coefficients D ₀, and apparent Young’s moduli E for substitutions onto the Opx M1 and M2 sites. Experimental r ₀ values for R³⁺ substitutions are 0.66–0.67 ? for M1 and 0.82–0.87 ? for M2. Simulations for enstatite result in r ₀ = 0.71–0.73 ? for M1 and ~0.79–0.87 ? for M2. Ferrosilite r ₀ values are systematically larger by ~0.05 ? for both M1 and M2. The latter is opposite to experimental literature data, which appear to show a slight decrease in $ r_{0}^{{{\text{M}}2}} $ r_{0}^{{{\text{M}}2}} in the presence of Fe. Additional systematic studies in Fe-bearing systems are required to resolve this inconsistency and to develop predictive Opx-melt partitioning models for use in terrestrial and lunar magmatic differentiation models.     

High-pressure electronic absorption spectroscopy of natural and synthetic Cr<Superscript>3+</Superscript>-bearing clinopyroxenes

Comparison of polarized optical absorption spectra of natural Ca-rich diopsides and synthetic NaCrSi2O6 and LiCrSi2O6 clinopyroxenes, evidences as vivid similarities, as noticeable differences. The similarities reflect the fact that in all cases Cr3+ enters the small octahedral M1-site of the clinopyroxene structure. The differences are due to some iron content in the natural samples causing broad intense near infrared bands of electronic spin-allowed dd transitions of Fe2+(M1, M2) and intervalence Fe2+/Fe3+ charge-transfer transition, and by different symmetry and different local crystal fields strength of Cr3+ in the crystal structures. The positions of the spin-allowed bands of Cr3+, especially of the low energy one caused by the electronic 4 A 2g → 2 T 1g transition, are found to be in accordance with mean M1–O distances. The local relaxation parameter ε calculated for limCr 3+ → 0 from the spectra and interatomic á Cr - O ñ \left\langle {Cr - O} \right\rangle and á Mg - O ñ \left\langle {Mg - O} \right\rangle distances yields a very high value, 0.96, indicating that in the clinopyroxene structure the local lattice relaxation around the “guest” ion, Cr3+, deviates greatly from the “diffraction” value, ε = 0, than in any other known Cr3+-bearing systems studied so far. Under pressure the spin-allowed bands of Cr3+ shift to higher energies and decrease in intensity quite in accordance with the crystal field theoretical expectations, while the spin-forbidden absorption lines remain practically unshifted, but also undergo a strong weakening. There is no evident dependence of the Racah parameter B of Cr3+ reflecting the covalence of the oxygen-chromium bond under pressure: within the uncertainty of determination it may be regarded as practically constant. The values of CrO6 octahedral modulus, k\textpoly\textloc k_{\text{poly}}^{\text{loc}} , derived from high-pressure spectra of natural chromium diopside and synthetic NaCrSi2O6 kosmochlor are very close, ~203 and ~196 GPa, respectively, being, however, nearly twice higher than that of MgO6 octahedron in diopside, 105(4) GPa, obtained by Thompson and Downs (2008). Such a strong stiffening of the structural octahedron, i.e. twice higher value of k\textCr3 + \textloc k_{{{\text{Cr}}^{3 + } }}^{\text{loc}} comparing with that of k\textMg2 + \textloc k_{{{\text{Mg}}^{2 + } }}^{\text{loc}} , may be caused by simultaneous substitution of Ca2+ by larger Na+ in the neighboring M2 sites at so-called jadeite-coupled substitution Mg2+ + Ca2+ → Cr3+ + Na+. It is also remarkable that the values of CrO6 octahedral modulus of NaCrSi2O6 kosmochlor obtained here are nearly twice larger than that of 90(16) GPa, evaluated by high-pressure X-ray structural refinement by Origlieri et al. (2003). Taking into account that the overall compressibility of the clinopyroxene structure should mainly be due to the compressibility of M1- and M2-sites, our k\textCr3 + \textloc k_{{{\text{Cr}}^{3 + } }}^{\text{loc}} -value, ~196 GPa, looks much more consistent with the bulk modulus value, 134(1) GPa.     

Crystal field and covalency of octahedral chromium in natural [8](Mg1?xCax) 3[6] (Al0.67Cr0.33)2Si3O12 garnets from upper mantle rocks

In the course of a thorough study of the influences of the second coordination sphere on the crystal field parameters of the 3d ^N -ions and the character of 3d ^N –O bonds in oxygen based minerals, 19 natural Cr³⁺-bearing (Mg,Ca)-garnets from upper mantle rocks were analysed and studied by electronic absorption spectroscopy, EAS. The garnets had compositions with populations of the ^[8] X-sites by 0.881 ± 0.053 (Ca + Mg) and changing Ca-fractions in the range 0.020 ≤ w _Ca[8] ≤ 0.745, while the ^[6] Y-site fraction was constant with x _Cr³⁺ _[6] = 0.335 ± 0.023. The garnets had colours from deeply violet-red for low Ca-contents (up to x _Ca = 0.28), grey with 0.28 ≤ x _Ca ≤ 0.4 and green with 0.4 ≤ x _Ca. The crystal field parameter of octahedral Cr³⁺ 10Dq decreases strongly on increasing Ca-fraction from 17,850 cm⁻¹ at x _Ca[8] = 0.020 to 16,580 cm⁻¹ at x _Ca[8] = 0.745. The data could be fit with two model which do statistically not differ: (1) two linear functions with a discontinuity close to x _Ca[8] ≈ 0.3,

A 57Fe M?ssbauer spectroscopic study of sogdianite: an example of a symmetric electric field gradient around Fe3+

Sogdianite, a double-ring silicate of composition ( \textZr_{0. 7 6} \textTi_{0. 3 8}^{4 +} \textFe_{0. 7 3}^{3 +} \textAl_0.13 )_{\Upsigma = 2} ( \square _{1. 1 5} \textNa_{0. 8 5} )_{\Upsigma = 2} \textK[\textLi ₃ \textSi _{1 2} \textO ₃₀ ] ( {\text{Zr}}_{0. 7 6} {\text{Ti}}_{0. 3 8}^{4 + } {\text{Fe}}_{0. 7 3}^{3 + } {\text{Al}}_{0.13} )_{\Upsigma = 2} \left( {\square_{ 1. 1 5} {\text{Na}}_{0. 8 5} } \right)_{\Upsigma = 2} {\text{K}}[{\text{Li}}_{ 3} {\text{Si}}_{ 1 2} {\text{O}}_{ 30} ] from Dara-i-Pioz, Tadjikistan, was studied by the combined application of ⁵⁷Fe M?ssbauer spectroscopy and electronic structure calculations. The M?ssbauer spectrum confirms published microprobe and X-ray single-crystal diffraction results that indicate that Fe³⁺ is located at the octahedral A-site and that no Fe²⁺ is present. Both the measured and calculated quadrupole splitting, ΔE _Q, for Fe³⁺ are virtually 0 mm s⁻¹. Such a value is unusually small for a silicate and it is the same as the ΔE _Q value for Fe³⁺ in structurally related sugilite. This result is traced back to the nearly regular octahedral coordination geometry corresponding to a very symmetric electric field gradient around Fe³⁺. A crystal chemical interpretation for the regular octahedral geometry and the resulting low ΔE _Q value for Fe³⁺ in the M?ssbauer spectrum of sogdianite is that structural strain is largely “taken up” by weak Li–O bonds permitting highly distorted LiO₄ tetrahedra. Weak Li–O bonding allows the edge-shared more strongly bonded Fe³⁺O₆ octahedra to remain regular in geometry. This may be a typical property for all double-ring silicates with tetrahedrally coordinated Li.     

Melting of carbonated pelites at 8–13 GPa: generating K-rich carbonatites for mantle metasomatism

The melting behaviour of three carbonated pelites containing 0–1 wt% water was studied at 8 and 13 GPa, 900–1,850°C to define conditions of melting, melt compositions and melting reactions. At 8 GPa, the fluid-absent and dry carbonated pelite solidi locate at 950 and 1,075°C, respectively; >100°C lower than in carbonated basalts and 150–300°C lower than the mantle adiabat. From 8 to 13 GPa, the fluid-present and dry solidi temperatures then increase to 1,150 and 1,325°C for the 1.1 wt% H₂O and the dry composition, respectively. The melting behaviour in the 1.1 wt% H₂O composition changes from fluid-absent at 8 GPa to fluid-present at 13 GPa with the pressure breakdown of phengite and the absence of other hydrous minerals. Melting reactions are controlled by carbonates, and the potassium and hydrous phases present in the subsolidus. The first melts, which composition has been determined by reverse sandwich experiments, are potassium-rich Ca–Fe–Mg-carbonatites, with extreme K₂O/Na₂O wt ratios of up to 42 at 8 GPa. Na is compatible in clinopyroxene with D_\textNa^{\textcpx/\textcarbonatite} = 10-18 D_{\text{Na}}^{{{\text{cpx}}/{\text{carbonatite}}}} = 10{-}18 at the solidus at 8 GPa. The melt K₂O/Na₂O slightly decreases with increasing temperature and degree of melting but strongly decreases from 8 to 13 GPa when K-hollandite extends its stability field to 200°C above the solidus. The compositional array of the sediment-derived carbonatites is congruent with alkali- and CO₂-rich melt or fluid inclusions found in diamonds. The fluid-absent melting of carbonated pelites at 8 GPa contrasts that at ≤5 GPa where silicate melts form at lower temperatures than carbonatites. Comparison of our melting temperatures with typical subduction and mantle geotherms shows that melting of carbonated pelites to 400-km depth is only feasible for extremely hot subduction. Nevertheless, melting may occur when subduction slows down or stops and thermal relaxation sets in. Our experiments show that CO₂-metasomatism originating from subducted crust is intimately linked with K-metasomatism at depth of >200 km. As long as the mantle remains adiabatic, low-viscosity carbonatites will rise into the mantle and percolate upwards. In cold subcontinental lithospheric mantle keels, the potassic Ca–Fe–Mg-carbonatites may freeze when reacting with the surrounding mantle leading to potassium-, carbonate/diamond- and incompatible element enriched metasomatized zones, which are most likely at the origin of ultrapotassic magmas such as group II kimberlites.     

Paragenesis of unusual Fe-cordierite (sekaninaite)-bearing paralava and clinker from the Kuznetsk coal basin,Siberia, Russia

Sekaninaite (X_Fe > 0.5)-bearing paralava and clinker are the products of ancient combustion metamorphism in the western part of the Kuznetsk coal basin, Siberia. The combustion metamorphic rocks typically occur as clinker beds and breccias consisting of vitrified sandstone–siltstone clinker fragments cemented by paralava, resulting from hanging-wall collapse above burning coal seams and quenching. Sekaninaite–Fe-cordierite (X_Fe = 95–45) is associated with tridymite, fayalite, magnetite, ± clinoferrosilite and ±mullite in paralava and with tridymite and mullite in clinker. Unmelted grains of detrital quartz occur in both rocks (<3 vol% in paralavas and up to 30 vol% in some clinkers). Compositionally variable siliceous, K-rich peraluminous glass is <30% in paralavas and up to 85% in clinkers. The paralavas resulted from extensive fusion of sandstone–siltstone (clinker), and sideritic/Fe-hydroxide material contained within them, with the proportion of clastic sediments ≫ ferruginous component. Calculated dry liquidus temperatures of the paralavas are 1,120–1,050°C and 920–1,050°C for clinkers, with calculated viscosities at liquidus temperatures of 10^1.6–7.0 and 10^7.0–9.8 Pa s, respectively. Dry liquidus temperatures of glass compositions range between 920 and 1,120°C (paralava) and 920–960°C (clinker), and viscosities at these temperatures are 10^9.7–5.5 and 10^8.8–9.7 Pa s, respectively. Compared with worldwide occurrences of cordierite–sekaninaite in pyrometamorphic rocks, sekaninaite occurs in rocks with X_Fe (mol% FeO/(FeO + MgO)) > 0.8; sekaninaite and Fe-cordierite occur in rocks with X_Fe 0.6–0.8, and cordierite (X_Fe < 0.5) is restricted to rocks with X_Fe < 0.6. The crystal-chemical formula of an anhydrous sekaninaite based on the refined structure is | \textK_0.02 |(\textFe_1.54^{2 +} \textMg_0.40 \textMn_0.06 )_{\Upsigma 2.00}^M [(\textAl_1.98 \textFe_0.02^{2 +} \textSi_1.00 )_{\Upsigma 3.00}^T1 (\textSi_3.94 \textAl_2.04 \textFe_0.02^{2 +} )_{\Upsigma 6.00}^T2 \textO₁₈ ]. \left| {{\text{K}}_{0.02} } \right|({\text{Fe}}_{1.54}^{2 + } {\text{Mg}}_{0.40} {\text{Mn}}_{0.06} )_{\Upsigma 2.00}^{M} [({\text{Al}}_{1.98} {\text{Fe}}_{0.02}^{2 + } {\text{Si}}_{1.00} )_{\Upsigma 3.00}^{T1} ({\text{Si}}_{3.94} {\text{Al}}_{2.04} {\text{Fe}}_{0.02}^{2 + } )_{\Upsigma 6.00}^{T2} {\text{O}}_{18} ].     

The magnesiowüstite: iron equilibrium and its implications for the activity-composition relations of (Mg,Fe)<Subscript>2</Subscript>SiO<Subscript>4</Subscript> olivine solid solutions

The chemical potential of oxygen (µO₂) in equilibrium with magnesiowüstite solid solution (Mg, Fe)O and metallic Fe has been determined by gas-mixing experiments at 1,473 K supplemented by solid-cell EMF experiments at lower temperatures. The results give:

Influence of temperature,composition, silica activity and oxygen fugacity on the H<Subscript>2</Subscript>O storage capacity of olivine at 8 GPa

Olivine crystals were grown in the presence of a hydrous silicate fluid during multi-anvil experiments at 8 GPa and 1,000–1,600°C. Experiments were conducted both in a simple system (FeO–MgO–SiO₂–H₂O) and in a more complex system containing additional elements (CaO–Na₂O–Al₂O₃–Cr₂O₃–TiO₂–FeO–MgO–SiO₂–H₂O). Silica activity was buffered by the presence of either pyroxene (high a _SiO2) or ferropericlase (low a _SiO2), and was buffered by the presence of Ni + NiO or Fe + FeO, or constrained by the presence of Fe₂O₃. Raman spectroscopy was used to identify pyroxene polymorphs in the run products. Clinoenstatite was present in the 1,000°C experiment, and enstatite in experiments at 1,400–1,520°C. The H₂O content of olivine was measured using secondary ion mass spectroscopy, and infrared spectroscopy was used to investigate the nature of hydrous defects. The H₂O storage capacity of olivine decreases with increasing temperature at 8 GPa. In contrast to previous experimental results at ≤2 GPa, no significant effect of varying oxygen fugacity is evident, but H₂O storage capacity is enhanced under conditions of low silica activity. No significant growth of low wavenumber (<3,400 cm⁻¹) peaks, generally associated with high at low pressure, was observed in the FTIR spectra of olivine from the high experiments. Our experiments show that previous high pressure H₂O storage capacity measurements for olivine synthesized under more oxidizing conditions than the Earth’s mantle are not likely to be compromised by the of the experiments. However, the considerable effect of temperature on H₂O storage capacity in olivine must be taken into account to avoid overestimation of the bulk upper mantle H₂O storage capacity.     

Quantifying strain birefringence halos around inclusions in diamond

The pressure and temperature conditions of formation of natural diamond can be estimated by measuring the residual stress that an inclusion remains under within a diamond. Raman spectroscopy has been the most commonly used technique for determining this stress by utilising pressure-sensitive peak shifts in the Raman spectrum of both the inclusion and the diamond host. Here, we present a new approach to measure the residual stress using quantitative analysis of the birefringence induced in the diamond. As the analysis of stress-induced birefringence is very different from that of normal birefringence, an analytical model is developed that relates the spherical inclusion size, R _i, host diamond thickness, L, and measured value of birefringence at the edge of the inclusion, \Updelta n(R_\texti )_\textav \Updelta n(R_{\text{i}} )_{\text{av}} , to the peak value of birefringence that has been encountered; to first order \Updelta n_\textpk = (3/4)(L/R_\texti )  \Updelta n(R_\texti )_\textav \Updelta n_{\text{pk}} = (3/4)(L/R_{\text{i}} ) \, \Updelta n(R_{\text{i}} )_{\text{av}} . From this birefringence, the remnant pressure (P _i) can be calculated using the photoelastic relationship \Updelta n_\textpk = - (3/4)n³ q_\textiso P_\texti \Updelta n_{\text{pk}} = - (3/4)n^{3} q_{\text{iso}} P_{\text{i}} , where q _iso is a piezo-optical coefficient, which can be assumed to be independent of crystallographic orientation, and n is the refractive index of the diamond. This model has been used in combination with quantitative birefringence analysis with a MetriPol system and compared to the results from both Raman point and 2D mapping analysis for a garnet inclusion in a diamond from the Udachnaya mine (Russia) and coesite inclusions in a diamond from the Finsch mine (South Africa). The birefringence model and analysis gave a remnant pressure of 0.53 ± 0.01 GPa for the garnet inclusion, from which a source pressure was calculated as 5.7 GPa at 1,175°C (temperature obtained from IR analysis of the diamond host). The Raman techniques could not be applied quantitatively to this sample to support the birefringence model; they were, however, applied to the largest coesite inclusion in the Finsch sample. The remnant pressure values obtained were 2.5 ± 0.1 GPa (birefringence), 2.5 ± 0.3 GPa (2D Raman map), and 2.5–2.6 GPa (Raman point analysis from all four inclusions). However, although the remnant pressures from the three methods were self-consistent, they led to anomalously low source pressure of 2.9 GPa at 1,150°C (temperature obtained from IR analysis) raising serious concerns about the use of the coesite-in-diamond geobarometer.     

Experimental determination of liquidus H<Subscript>2</Subscript>O contents of haplogranite at deep-crustal conditions

The liquidus water content of a haplogranite melt at high pressure (P) and temperature (T) is important, because it is a key parameter for constraining the volume of granite that could be produced by melting of the deep crust. Previous estimates based on melting experiments at low P (≤0.5 GPa) show substantial scatter when extrapolated to deep crustal P and T (700–1000 °C, 0.6–1.5 GPa). To improve the high-P constraints on H₂O concentration at the granite liquidus, we performed experiments in a piston–cylinder apparatus at 1.0 GPa using a range of haplogranite compositions in the albite (Ab: NaAlSi₃O₈)—orthoclase (Or: KAlSi₃O₈)—quartz (Qz: SiO₂)—H₂O system. We used equal weight fractions of the feldspar components and varied the Qz between 20 and 30 wt%. In each experiment, synthetic granitic composition glass + H₂O was homogenized well above the liquidus T, and T was lowered by increments until quartz and alkali feldspar crystalized from the liquid. To establish reversed equilibrium, we crystallized the homogenized melt at the lower T and then raised T until we found that the crystalline phases were completely resorbed into the liquid. The reversed liquidus minimum temperatures at 3.0, 4.1, 5.8, 8.0, and 12.0 wt% H₂O are 935–985, 875–900, 775–800, 725–775, and 650–675 °C, respectively. Quenched charges were analyzed by petrographic microscope, scanning electron microscope (SEM), X-ray diffraction (XRD), and electron microprobe analysis (EMPA). The equation for the reversed haplogranite liquidus minimum curve for Ab_36.25Or_36.25Qz_27.5 (wt% basis) at 1.0 GPa is \(T = - 0.0995 w_{{{\text{H}}_{ 2} {\text{O}}}}^{ 3} + 5.0242w_{{{\text{H}}_{ 2} {\text{O}}}}^{ 2} - 88.183 w_{{{\text{H}}_{ 2} {\text{O}}}} + 1171.0\) for \(0 \le w_{{{\text{H}}_{ 2} {\text{O}}}} \le 17\) wt% and \(T\) is in °C. We present a revised \(P - T\) diagram of liquidus minimum H₂O isopleths which integrates data from previous determinations of vapor-saturated melting and the lower pressure vapor-undersaturated melting studies conducted by other workers on the haplogranite system. For lower H₂O (<5.8 wt%) and higher temperature, our results plot on the high end of the extrapolated water contents at liquidus minima when compared to the previous estimates. As a consequence, amounts of metaluminous granites that can be produced from lower crustal biotite–amphibole gneisses by dehydration melting are more restricted than previously thought.     

Structural relaxation and crystal field stabilization in Cr3+-containing oxides and silicates

The effect of crystal structure relaxation in oxygen-based Cr³⁺-containing minerals on the crystal field stabilization energy (CFSE) is considered. It is shown that the dependence of \textCFSE_{\textCr ³⁺} {\text{CFSE}}_{{{\text{Cr}}^{ 3+ } }} , which is found from optical absorption spectra, on the average interatomic distances is described by the power function with a negative exponent c \mathord