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1.
Sogdianite, a double-ring silicate of composition
( \text Zr0. 7 6 \text Ti0. 3 84 + \text Fe0. 7 33 + \text Al0.13 ) \Upsigma = 2 ( \square 1. 1 5 \text Na0. 8 5 ) \Upsigma = 2 \text K[\text Li 3 \text Si 1 2 \text O 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 Fe 3+ is located at the octahedral A-site and that no Fe 2+ is present. Both the measured and calculated quadrupole splitting, Δ E
Q, for Fe 3+ are virtually 0 mm s −1. Such a value is unusually small for a silicate and it is the same as the Δ E
Q value for Fe 3+ 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 Fe 3+. A crystal chemical interpretation for the regular octahedral geometry and the resulting low Δ E
Q value for Fe 3+ in the M?ssbauer spectrum of sogdianite is that structural strain is largely “taken up” by weak Li–O bonds permitting highly
distorted LiO 4 tetrahedra. Weak Li–O bonding allows the edge-shared more strongly bonded Fe 3+O 6 octahedra to remain regular in geometry. This may be a typical property for all double-ring silicates with tetrahedrally
coordinated Li. 相似文献
2.
Monticellite is a common magmatic mineral in the groundmass of kimberlites. A new oxygen barometer for kimberlite magmas is
calibrated based on the Fe content of monticellite, CaMgSiO 4, in equilibrium with kimberlite liquids in experiments at 100 kPa from 1,230 to 1,350°C and at log fO 2 from NNO-4.1 to NNO+5.3 (where NNO is the nickel–nickel oxide buffer). The XFe Mtc/ XFe liq was found to decrease with increasing fO 2, consistent with only Fe 2+ entering the monticellite structure. Although the XFe-in-monticellite varies with temperature and composition, these dependencies are small compared to that with fO 2. The experimental data were fitted by weighted least square regression to the following relationship:
\Updelta \text NNO = \frac{ log[ 0.858( ±0.021)\frac X\textFe\textLiq X\textFe\textMtc ] - 0.139( ±0.022) }0.193( ±0.004) \Updelta {\text{NNO}} = \frac{{\left\{ {\log \left[ {0.858( \pm 0.021)\frac{{X_{\text{Fe}}^{\text{Liq}} }}{{X_{\text{Fe}}^{\text{Mtc}} }}} \right] - 0.139( \pm 0.022)} \right\}}}{0.193( \pm 0.004)} where ΔNNO is the fO 2 relative to that of the Nickel-bunsenite (NNO) buffer and XFe liq/ XFe Mtc is the ratio of mole fraction of Fe in liquid and Fe-in-monticellite (uncertainties at 2σ). The application of this oxygen
barometer to natural kimberlites from both the literature and our own investigations, assuming the bulk rock FeO is that of
their liquid FeO, revealed a range in fO 2 from NNO-3.5 to NNO+1.7. A range of Mg/(Mg + Fe 2+) (Mg#) for kimberlite melts of 0.46–0.88 was derived from the application of the experimentally determined monticellite-liquid
Kd Fe 2+–Mg to natural monticellites. The range in Mg# is broader and less ultramafic than previous estimates of kimberlites, suggesting
an evolution under a wide range of petrologic conditions. 相似文献
3.
The onset of hydrous partial melting in the mantle above the transition zone is dictated by the H 2O 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. H 2O 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 H 2O 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 H 2O 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 ( \text ppm ) = 57.6( ±16 ) × P( \text GPa ) - 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 H 2O storage capacity, which is 440 ± 200 ppm at 400 km. This suggests that MORB source upper mantle, which contains 50–200 ppm
bulk H 2O, 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 H 2O, can exceed the peridotite H 2O 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 H 2O 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
4.
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 fO 2 to be derived:
\text H 2 \text O( \text wt% ) = \text H 2 \text O\textMORB ( \text wt% ) + ( 5.84 ×10 - 5 *\text P - 2.29 ×10 - 2 ) ×( \text Na2 \text O + \text K2 \text O )( \text wt% ) + 4.67 ×10 - 2 ×\Updelta \text NNO - 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 H 2O 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 H 2O solubility data on basaltic melts at pressures above 1 kbar yields a value of 19 cm 3/mol in reasonable agreement with estimates obtained from density measurements. 相似文献
5.
The present work aims in discussing a principle that distinguishes between elastic parameters sets, $ \{ \Upphi \} \equiv \{ K_{0} , \, K^{\prime}, \, V_{0} ,\ldots\} The present work aims in discussing a principle that distinguishes between elastic parameters sets,
{ \Upphi } o { K0 , K￠, V0 ,?} \{ \Upphi \} \equiv \{ K_{0} , \, K^{\prime}, \, V_{0} ,\ldots\} , on the basis of an energetic criterion: once a reference set,
{ \UpphiR } \{ \Upphi_{R} \} , is given, another one can be fixed,
{ \Upphi min } \left\{ {\Upphi_{ \min } } \right\} , so that they are as close as possible to each other, but yield non-equivalent deformation energy curves
\Updelta G({ \Upphi } )\textdeform \Updelta G(\{ \Upphi \} )_{\text{deform}} , i.e. they give
\Updelta G({ \UpphiR } )\textdeform \Updelta G(\{ \Upphi_{R} \} )_{\text{deform}} and
\Updelta G({ \Upphi min } )\textdeform \Updelta G(\{ \Upphi_{ \min } \} )_{\text{deform}} such that
| \Updelta G({ \Upphi min } )\textdeform - \Updelta G({ \UpphiR } )\textdeform | 3 1×s[\Updelta G\textdeform ]. \left| {\Updelta G(\{ \Upphi_{ \min } \} )_{\text{deform}} - \Updelta G(\{ \Upphi_{R} \} )_{\text{deform}} } \right| \ge 1\times \sigma [\Updelta G_{\text{deform}} ]. ΔG
deform, calculated using the equation of state (EoS), and its uncertainty σ[ΔG
deform], obtained by a propagation of the errors affecting
{ \Upphi } \{ \Upphi \} are crucial to fix which mineral assemblage forms at P–T conditions and allow one to assess the reliability of such a prediction. We explore some properties related to the principle
introduced, using the average values of the elastic parameters found in literature and related uncertainties for di-octahedral
mica, olivine, garnet and clinopyroxene. Two elementary applications are briefly discussed: the effect of refining V
0 in fitting EoSs to P–V experimental data, in the case of garnet and omphacite, and the phengite 3T–2M
1 relative stability, controlled by pressure. 相似文献
6.
Diffusion of Li under anhydrous conditions at 1 atm and under fluid-present elevated pressure (1.0–1.2 GPa) conditions has
been measured in natural zircon. The source of diffusant for 1-atm experiments was ground natural spodumene, which was sealed
under vacuum in silica glass capsules with polished slabs of zircon. An experiment using a Dy-bearing source was also conducted
to evaluate possible rate-limiting effects on Li diffusion of slow-diffusing REE +3 that might provide charge balance. Diffusion experiments performed in the presence of H 2O–CO 2 fluid were run in a piston–cylinder apparatus, using a source consisting of a powdered mixture of spodumene, quartz and zircon
with oxalic acid added to produce H 2O–CO 2 fluid. Nuclear reaction analysis (NRA) with the resonant nuclear reaction 7Li(p,γ) 8Be was used to measure diffusion profiles for the experiments. The following Arrhenius parameters were obtained for Li diffusion
normal to the c-axis over the temperature range 703–1.151°C at 1 atm for experiments run with the spodumene source:
|
D\textLi = 7.17 ×10 - 7 exp( - 275 ±11 \textkJmol - 1 /\textRT)\textm2 \texts - 1. D_{\text{Li}} = 7.17 \times 10^{ - 7} { \exp }( - 275 \pm 11\,{\text{kJmol}}^{ - 1} /{\text{RT}}){\text{m}}^{2} {\text{s}}^{ - 1}. 相似文献
7.
Mineral-specific IR absorption coefficients were calculated for natural and synthetic olivine, SiO 2 polymorphs, and GeO 2 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 SiO 2 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
\text l \text mol\textH2\textO-1 \text cm-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
\text l \text mol\textH2\textO-1 \text cm-2\text{l}\,\text{mol}_{{\text{H}_2}\text{O}}^{-1}\,\text{cm}^{-2}; stishovite ε = 485,000 ± 109,000
\text l \text mol\textH2\textO-1 \text cm-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
8.
Experiments were conducted to determine CO 2 solubilities in alkali basalts from Vesuvius, Etna and Stromboli volcanoes. The basaltic melts were equilibrated with nearly
pure CO 2 at 1,200°C under oxidizing conditions and at pressures ranging from 269 to 2,060 bars. CO 2 solubility was determined by FTIR measurements. The results show that alkalis have a strong effect on the CO 2 solubility and confirm and refine the relationship between the compositional parameter Π devised by Dixon (Am Mineral 82:368–378,
1997) and the CO 2 solubility. A general thermodynamic model for CO 2 solubility in basaltic melts is defined for pressures up to 2 kbars. Based on the assumption that O 2− and CO 32− mix ideally, we have:
|
_boxclose_3^2 - ^m (P,T)X_^2 - ^m f__2 (P,T) K(P,T) = X__3^2 - ^m (P,T) ( X_^2 - ^m f__2 (P,T) ). \begin{gathered} K(P,T) = {\frac{{X_{{{\text{CO}}_{3}^{2 - } }}^{m} (P,T)}}{{X_{{{\text{O}}^{2 - } }}^{m} \times f_{{{\text{CO}}_{2} }} (P,T)}}} \hfill \\ K(P,T) = {{X_{{{\text{CO}}_{3}^{2 - } }}^{m} (P,T)} \mathord{\left/ {\vphantom {{X_{{{\text{CO}}_{3}^{2 - } }}^{m} (P,T)} {\left( {X_{{{\text{O}}^{2 - } }}^{m} \times f_{{{\text{CO}}_{2} }} (P,T)} \right).}}} \right. \kern-\nulldelimiterspace} {\left( {X_{{{\text{O}}^{2 - } }}^{m} \times f_{{{\text{CO}}_{2} }} (P,T)} \right).}} \hfill \\ \end{gathered} 相似文献
9.
The addition of Fe and Cr to the simple system MgO-SiO 2-Al 2O 3 markedly affects the activities of phases involved in the equilibrium |
\textMg\text2 \textSiO\text4 \text + MgAl\text2 \textSiO\text6 \text = MgAl\text2 \textO\text4 \text + Mg\text2 \textSi\text2 \textO\text6 \textOlivine + Opx\textsolid solution \text = Spinel + Opx\textsolid solution \begin{gathered} {\text{Mg}}_{\text{2}} {\text{SiO}}_{\text{4}} {\text{ + MgAl}}_{\text{2}} {\text{SiO}}_{\text{6}} {\text{ = MgAl}}_{\text{2}} {\text{O}}_{\text{4}} {\text{ + Mg}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} \hfill \\ {\text{Olivine + Opx}}_{{\text{solid solution}}} {\text{ = Spinel + Opx}}_{{\text{solid solution}}} \hfill \\ \end{gathered} 相似文献
10.
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)] \textFe \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 Fe 2+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. 相似文献
11.
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 H 2O vibrational modes. The loss of H 2O molecules caused a sequence of structural transitions in which the host framework transformation path was coupled primarily
via the thermal motion of guest Na +/Ca 2+ cations and H 2O molecules. The observation of different interactions of H 2O molecules and Na +/Ca 2+ 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–H 2O 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–H 2O–Al–Si framework interactions in NAT-type channels, and these bands can be used to evaluate the energetic evolution of Na +/Ca 2+ 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 +/Ca 2+ order–disorder that occur when mesolite loses its Ow4 H 2O molecules. 相似文献
12.
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–Al 2O 3–SiO 2 and subsystem CaO–MgO–SiO 2. 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. 相似文献
13.
Uranium mineralization in the El Erediya area, Egyptian Eastern Desert, has been affected by both high temperature and low
temperature fluids. Mineralization is structurally controlled and is associated with jasperoid veins that are hosted by a
granitic pluton. This granite exhibits extensive alteration, including silicification, argillization, sericitization, chloritization,
carbonatization, and hematization. The primary uranium mineral is pitchblende, whereas uranpyrochlore, uranophane, kasolite,
and an unidentified hydrated uranium niobate mineral are the most abundant secondary uranium minerals. Uranpyrochlore and
the unidentified hydrated uranium niobate mineral are interpreted as alteration products of petscheckite. The chemical formula
of the uranpyrochlore based upon the Electron Probe Micro Analyzer (EPMA) is . It is characterized by a relatively high Zr content (average ZrO 2 = 6.6 wt%). The average composition of the unidentified hydrated uranium niobate mineral is , where U and Nb represent the dominant cations in the U and Nb site, respectively. Uranophane is the dominant U 6+ silicate phase in oxidized zones of the jasperoid veins. Kasolite is less abundant than uranophane and contains major U,
Pb, and Si but only minor Ca, Fe, P, and Zr. A two-stage metallogenetic model is proposed for the alteration processes and
uranium mineralization at El Erediya. The primary uranium minerals were formed during the first stage of the hydrothermal
activity that formed jasperoid veins in El Eradiya granite (130–160 Ma). This stage is related to the Late Jurassic–Early
Cretaceous phase of the final Pan-African tectono-thermal event in Egypt. After initial formation of El Erediya jasperoid
veins, a late stage of hydrothermal alteration includes argillization, dissolution of iron-bearing sulfide minerals, formation
of iron-oxy hydroxides, and corrosion of primary uranium minerals, resulting in enrichment of U, Ca, Pb, Zr, and Si. During
this stage, petscheckite was altered to uranpyrochlore and oxy-petscheckite. Uranium was likely transported as uranyl carbonate
and uranyl fluoride complexes. With change of temperature and pH, these complexes became unstable and combined with silica,
calcium, and lead to form uranophane and kasolite. Finally, at a later stage of low-temperature supergene alteration, oxy-petscheckite
was altered to an unidentified hydrated uranium niobate mineral by removal of Fe. 相似文献
14.
To get deeper insight into the phase relations in the end-member system Fe 2SiO 4 and in the system (Fe, Mg) 2SiO 4 experiments were performed in a multi-anvil apparatus at 7 and 13 GPa and 1,000–1,200°C as a function of oxygen fugacity.
The oxygen fugacity was varied using the solid oxygen buffer systems Fe/FeO, quartz–fayalite–magnetite, MtW and Ni/NiO. The
run products were characterized by electron microprobe, Raman- and FTIR-spectroscopy, X-ray powder diffraction and transmission
electron microscopy. At fO 2 corresponding to Ni/NiO Fe-ringwoodite transforms to ferrosilite and spinelloid according to the reaction: 9 Fe 2SiO 4 + O 2 = 6 FeSiO 3 + 5 Fe 2.40Si 0.60O 4. Refinement of site occupancies in combination with stoichiometric Fe 3+ calculations show that 32% of the total Fe is incorporated as Fe 3+ according to From the Rietveld refinement we identified spl as spinelloid III (isostructural with wadsleyite) and/or spinelloid V. As
we used water in excess in the experiments the run products were also analyzed for structural water incorporation. Adding
Mg to the system increases the stability field of ringwoodite to higher oxygen fugacity and the spinel structure seems to
accept higher Fe 3+ but also water concentrations that may be linked. At oxygen fugacity corresponding to MtW conditions similar phase relations
in respect to the breakdown reaction in the Fe-end-member system were observed but with a strong fractionation of Fe into
spl and Mg into coexisting cpx. Thus, through this strong fractionation it is possible to stabilize very Fe-rich wadsleyite
with considerable Fe 3+ concentrations even at an intermediate Fe–Mg bulk composition: assuming constant K
D independent on composition and a bulk composition of x
Fe = 0.44 this fractionation would stabilize spl with x
Fe = 0.72. Thus, spl could be a potential Fe 3+ bearing phase at P–T conditions of the transition zone but because of the oxidizing conditions and the Fe-rich bulk composition
needed one would expect it more in subduction zone environments than in the transition zone in senso stricto.
相似文献
15.
A natural Ca-poor pigeonite (Wo 6En 76Fs 18) from the ureilite meteorite sample PCA82506-3, free of exsolved augite, was studied by in situ high-temperature single-crystal
X-ray diffraction. The sample, monoclinic P2 1/ c, was annealed up to 1,093°C to induce a phase transition from P2 1/ c to C2/ c symmetry. The variation with increasing temperature of the lattice parameters and of the intensity of the b-type reflections ( h + k = 2 n + 1, present only in the P2 1/ c phase) showed a displacive phase transition P2 1/ c to C2/ c at a transition temperature T
Tr = 944°C, first order in character. The Fe–Mg exchange kinetics was studied by ex situ single-crystal X-ray diffraction in
a range of temperatures between the closure temperature of the Fe–Mg exchange reaction and the transition temperature. Isothermal
disordering annealing experiments, using the IW buffer, were performed on three crystals at 790, 840 and 865°C. Linear regression
of ln k
D versus 1/ T yielded the following equation:
ln k\textD = - 3717( ±416)/ T( K) + 1.290( ±0.378); ( R2 = 0.988) \ln \,k_{\text{D}} = - 3717( \pm 416)/T(K) + 1.290( \pm 0.378);\quad (R^{2} = 0.988) . The closure temperature ( T
c) calculated using this equation was ∼740(±30)°C. Analysis of the kinetic data carried out taking into account the e.s.d.'s
of the atomic fractions used to define the Fe–Mg degree of order, performed according to Mueller’s model, allowed us to retrieve
the disordering rate constants C
0
K
dis+ for all three temperatures yielding the following Arrhenius relation:
ln( C0 K\textdis + ) = ln K0 - Q/( RT) = 20.99( ±3.74) - 26406( ±4165)/ T( K); ( R2 = 0.988) \ln \left( {C_{0} K_{\text{dis}}^{ + } } \right) = \ln \,K_{0} - Q/(RT) = 20.99( \pm 3.74) - 26406( \pm 4165)/T(K);\quad (R^{2} = 0.988) . An activation energy of 52.5(±4) kcal/mol for the Fe–Mg exchange process was obtained. The above relation was used to calculate
the following Arrhenius relation modified as a function of X
Fe (in the range of X
Fe = 0.20–0.50):
ln( C0 K\textdis + ) = (21.185 - 1.47 X\textFe ) - \frac(27267 - 4170 X\textFe ) T( K) \ln \left( {C_{0} K_{\text{dis}}^{ + } } \right) = (21.185 - 1.47X_{\text{Fe}} ) - {\frac{{(27267 - 4170X_{\text{Fe}} )}}{T(K)}} . The cooling time constant, η = 6 × 10 −1 K −1 year −1 calculated on the PCA82506-3 sample, provided a cooling rate of the order of 1°C/min consistent with the extremely fast late
cooling history of the ureilite parent body after impact excavation. 相似文献
16.
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 | / |
\vphantom c [`(R)]n [`(R)]n {c \mathord{\left/ {\vphantom {c {\bar{R}^{n} }}} \right. \kern-\nulldelimiterspace} {\bar{R}^{n} }} , where n approaches 5, as predicted theoretically, for pure Cr3+ compounds, but decreases to 1.0–1.5 for Cr3+-containing oxide and silicate solid solutions. The deviation of the experimental dependence for solid solutions from the
theoretical curve is due to structure relaxation, which tends to bring the local structure of Cr3+ ions closer to the structure in the pure Cr compound, thus producing changes in interatomic distances between the nearest
neighbors with respect to those in the average structure determined by X-ray diffraction. As a consequence, the mixing enthalpy
of Cr3+-bearing solid solutions can be represented by the sum of contributions from lattice strain and CFSE. The latter contribution
is most often negative in sign and, therefore, brings the Al–Cr solid solutions close to an ideal solid solution. It is supposed
that the increased Cr content in minerals from deep-seated mantle xenoliths and mineral inclusions in diamonds results from
the effect of
\textCFSE\textCr 3+ {\text{CFSE}}_{{{\text{Cr}}^{ 3+ } }} enhanced by high pressure. 相似文献
17.
Three Al-Cr exchange isotherms at 1,250°, 1,050°, and 796° between Mg(Al, Cr) 2O 4 spinel and (Al, Cr) 2O 3 corundum crystalline solutions have been studied experimentally at 25 kbar pressure. Starting from gels of suitable bulk
compositions, close approach to equilibrium has been demonstrated in each case by time studies.
Using the equation of state for (Al, Cr) 2O 3 crystalline solution (Chatterjee et al. 1982a) and assuming that the Mg(Al, Cr) 2O 4 can be treated in terms of the asymmetric Margules relation, the exchange isotherms were solved for Δ G
*,
and
. The best constrained data set from the 1,250° C isotherm clearly shows that the latter two quantities do not overlap within
three standard deviations, justifying the choice of asymmetric Margules relation for describing the excess mixing properties
of Mg(Al, Cr) 2O 4 spinels. Based on these experiments, the following polybaric-polythermal equation of state can be formulated:
, P expressed in bars, T in K, G
m
ex
and W
G,i
Sp
in joules/mol.
Temperature-dependence of G
m
ex
is best constrained in the range 796–1,250° C; extrapolation beyond that range would have to be done with caution. Such extrapolation
to lower temperature shows tentatively that at 1 bar pressure the critical temperature, T
c, of the spinel solvus is 427° C, with dT c/dP≈1.3 K/kbar. The critical composition, X
c, is 0.42
, and changes barely with pressure.
Substantial error in calculated phase diagrams will result if the significant positive deviation from ideality is ignored
for Al-Cr mixing in such spinels. 相似文献
18.
Comparison of polarized optical absorption spectra of natural Ca-rich diopsides and synthetic NaCrSi 2O 6 and LiCrSi 2O 6 clinopyroxenes, evidences as vivid similarities, as noticeable differences. The similarities reflect the fact that in all
cases Cr 3+ 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 Fe 2+(M1, M2) and intervalence Fe 2+/Fe 3+ charge-transfer transition, and by different symmetry and different local crystal fields strength of Cr 3+ in the crystal structures. The positions of the spin-allowed bands of Cr 3+, 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 lim Cr
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, Cr 3+, deviates greatly from the “diffraction” value, ε = 0, than in any other known Cr 3+-bearing systems studied so far. Under pressure the spin-allowed bands of Cr 3+ 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 Cr 3+ 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 CrO 6 octahedral modulus,
k\textpoly\textloc k_{\text{poly}}^{\text{loc}} , derived from high-pressure spectra of natural chromium diopside and synthetic NaCrSi 2O 6 kosmochlor are very close, ~203 and ~196 GPa, respectively, being, however, nearly twice higher than that of MgO 6 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 Ca 2+ by larger Na + in the neighboring M2 sites at so-called jadeite-coupled substitution Mg 2+ + Ca 2+ → Cr 3+ + Na +. It is also remarkable that the values of CrO 6 octahedral modulus of NaCrSi 2O 6 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. 相似文献
19.
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 (HClO 4–HCl and NaCl–NaClO 4 mixtures, and solutions of MgCl 2 and CaCl 2). Specific interaction theory analysis of these formation constants produced coherent estimates of (a) PbCl +,
\text PbCl20 {\text{PbCl}}_{2}^{0} , and PbCl 3− 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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log \text Cl b 1 = 1. 4 9 1- 2.0 4 I 1/ 2 ( 1+ 1. 5 I 1/ 2 ) - 1 + 0. 2 3 8 I log \text Cl b 2 = 2.0 6 2- 3.0 6 I 1/ 2 ( 1+ 1. 5 I 1/ 2 ) - 1 + 0. 3 6 9 I log \text Cl b 3 = 1. 8 9 9- 3.0 6 I 1/ 2 ( 1+ 1. 5 I 1/ 2 ) - 1 + 0. 4 3 9 I. \begin{gathered} {\log}\,{}_{\text{ Cl}} \beta_{ 1} = 1. 4 9 1- 2.0 4\,I^{ 1/ 2} \left( { 1+ 1. 5\,I^{ 1/ 2} } \right)^{ - 1} +\,0. 2 3 8\,I \hfill \\ {\log}\,{}_{\text{ Cl}} \beta_{ 2} = 2.0 6 2- 3.0 6\,I^{ 1/ 2} \left( { 1+ 1. 5\,I^{ 1/ 2} } \right)^{ - 1} +\,0. 3 6 9\,I \hfill \\ {\log}\,{}_{\text{ Cl}} \beta_{ 3} = 1. 8 9 9- 3.0 6\,I^{ 1/ 2} \left( { 1+ 1. 5\,I^{ 1/ 2} } \right)^{ - 1} +\,0. 4 3 9\,I. \hfill \\ \end{gathered} 相似文献
20.
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):
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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
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