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1.
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. 相似文献
2.
Priscille Lesne Bruno Scaillet Michel Pichavant Giada Iacono-Marziano Jean-Michel Beny 《Contributions to Mineralogy and Petrology》2011,162(1):133-151
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 H2OMORB 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. 相似文献
3.
David L. Naftz Frank J. Millero Blair F. Jones W. Reed Green 《Aquatic Geochemistry》2011,17(6):809-820
Great Salt Lake (GSL) is one of the largest and most saline lakes in the world. In order to accurately model limnological
processes in GSL, hydrodynamic calculations require the precise estimation of water density (ρ) under a variety of environmental conditions. An equation of state was developed with water samples collected from GSL to
estimate density as a function of salinity and water temperature. The ρ of water samples from the south arm of GSL was measured as a function of temperature ranging from 278 to 323 degrees Kelvin
(oK) and conductivity salinities ranging from 23 to 182 g L−1 using an Anton Paar density meter. These results have been used to develop the following equation of state for GSL (σ = ± 0.32 kg m−3):
r- r0 = 184.0 10 6 2 + 1.0 4 70 8*\textS - 1. 2 10 6 1*\textT + 3. 1 4 7 2 1 \textE - 4*\textS 2 + 0.00 1 9 9 \textT 2 - 0.00 1 1 2*\textS*\textT, \rho - \rho^{0} = { 184}.0 10 6 2 { } + { 1}.0 4 70 8*{\text{S}} - 1. 2 10 6 1*{\text{T }} + { 3}. 1 4 7 2 1 {\text{E}} - 4*{\text{S}}^{ 2} + \, 0.00 1 9 9 {\text{T}}^{ 2} - 0.00 1 1 2*{\text{S}}*{\text{T}}, 相似文献
4.
The onset of hydrous partial melting in the mantle above the transition zone is dictated by the H2O storage capacity of peridotite, which is defined as the maximum concentration that the solid assemblage can store at P and T without stabilizing a hydrous fluid or melt. H2O storage capacities of minerals in simple systems do not adequately constrain the peridotite water storage capacity because
simpler systems do not account for enhanced hydrous melt stability and reduced H2O activity facilitated by the additional components of multiply saturated peridotite. In this study, we determine peridotite-saturated
olivine and pyroxene water storage capacities at 10–13 GPa and 1,350–1,450°C by employing layered experiments, in which the
bottom ~2/3 of the capsule consists of hydrated KLB-1 oxide analog peridotite and the top ~1/3 of the capsule is a nearly
monomineralic layer of hydrated Mg# 89.6 olivine. This method facilitates the growth of ~200-μm olivine crystals, as well
as accessory low-Ca pyroxenes up to ~50 μm in diameter. The presence of small amounts of hydrous melt ensures that crystalline
phases have maximal H2O contents possible, while in equilibrium with the full peridotite assemblage (melt + ol + pyx + gt). At 12 GPa, olivine and
pyroxene water storage capacities decrease from ~1,000 to 650 ppm, and ~1,400 to 1,100 ppm, respectively, as temperature increases
from 1,350 to 1,450°C. Combining our results with those from a companion study at 5–8 GPa (Ardia et al., in prep.) at 1,450°C,
the olivine water storage capacity increases linearly with increasing pressure and is defined by the relation
C\textH2 \textO\textolivine ( \textppm ) = 57.6( ±16 ) ×P( \textGPa ) - 169( ±18 ). C_{{{\text{H}}_{2} {\text{O}}}}^{\text{olivine}} \left( {\text{ppm}} \right) = 57.6\left( { \pm 16} \right) \times P\left( {\text{GPa}} \right) - 169\left( { \pm 18} \right). Adjustment of this trend for small increases in temperature along the mantle geotherm, combined with experimental determinations
of
D\textH2 \textO\textpyx/olivine D_{{{\text{H}}_{2} {\text{O}}}}^{\text{pyx/olivine}} from this study and estimates of
D\textH2 \textO\textgt/\textolivine D_{{{\text{H}}_{2} {\text{O}}}}^{{{\text{gt}}/{\text{olivine}}}} , allows for estimation of peridotite H2O storage capacity, which is 440 ± 200 ppm at 400 km. This suggests that MORB source upper mantle, which contains 50–200 ppm
bulk H2O, is not wet enough to incite a global melt layer above the 410-km discontinuity. However, OIB source mantle and residues
of subducted slabs, which contain 300–1,000 ppm bulk H2O, can exceed the peridotite H2O storage capacity and incite localized hydrous partial melting in the deep upper mantle. Experimentally determined values
of
D\textH2 \textO\textpyx/\textolivine D_{{{\text{H}}_{2} {\text{O}}}}^{{{\text{pyx}}/{\text{olivine}}}} at 10–13 GPa have a narrow range of 1.35 ± 0.13, meaning that olivine is probably the most important host of H2O in the deep upper mantle. The increase in hydration of olivine with depth in the upper mantle may have significant influence
on viscosity and other transport properties. 相似文献
5.
Lead chloride formation constants at 25°C were derived from analysis of previous spectrophotometrically generated observations
of lead speciation in a variety of aqueous solutions (HClO4–HCl and NaCl–NaClO4 mixtures, and solutions of MgCl2 and CaCl2). Specific interaction theory analysis of these formation constants produced coherent estimates of (a) PbCl+,
\textPbCl20 {\text{PbCl}}_{2}^{0} , and PbCl3− formation constants at zero ionic strength, and (b) well-defined depictions of the dependence of these formation constants
on ionic strength. Accompanying examination of a recent IUPAC critical assessment of lead formation constants, in conjunction
with the spectrophotometrically generated formation constants presented in this study, revealed significant differences among
various subsets of the IUPAC critically selected data. It was found that these differences could be substantially reduced
through reanalysis of the formation constant data of one of the subsets. The resulting revised lead chloride formation constants
are in good agreement with the formation constants derived from the earlier spectrophotometrically generated data. Combining
these data sets provides an improved characterization of lead chloride complexation over a wide range of ionic strengths:
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