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1.
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
2.
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
3.
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
4.
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
| \text K0.02 |(\text Fe1.542 + \text Mg0.40 \text Mn0.06 ) \Upsigma 2.00M [(\text Al1.98 \text Fe0.022 + \text Si1.00 ) \Upsigma 3.00T1 (\text Si3.94 \text Al2.04 \text Fe0.022 + ) \Upsigma 6.00T2 \text O18 ]. \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} ]. 相似文献
5.
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
6.
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
7.
Electron paramagnetic resonance (EPR) study of single crystals of chromium-doped forsterite grown by the Czochralski method
in two different research laboratories has revealed, apart from the known paramagnetic centers Cr 3+( M1), Cr 3+( M2) and Cr 4+, a new center
\text Cr 3+ ( M 1)- V\textMg 2+ ( M 2) {\text{Cr}}^{ 3+ } (M 1){-}V_{{{\text{Mg}}^{ 2+ } }} (M 2) formed by a Cr 3+ ion substituting for Mg 2+ at the M1 structural position with a nearest-neighbor Mg 2+ vacancy at the M2 position. For this center, the conventional zero-field splitting parameters D and E and the principal g values and A values of the 53Cr hyperfine splitting have been determined as follows: D = 33.95(3) GHz, E = 8.64(1) GHz, g = [1.9811(2), 1.9787(2), 1.9742(2)], A = [51(3), 52(2), 44(3)] MHz. The center has been identified by comparing EPR spectra with those of the charge-uncompensated
ion Cr 3+( M1) and the ion pair Cr 3+( M1)–Li +( M2) observed in forsterite crystals codoped with chromium and lithium. It has been found that the concentration of the new
center decreases to zero, whereas that of the Cr 3+( M1) and Cr 3+( M1)–Li +( M2) centers increases with an increase of the Li content from 0 up to ~0.03 wt% (at the same Cr content ~0.07 wt%) in the melt.
The known low-temperature luminescence data pertinent to the centers under consideration are also discussed. 相似文献
8.
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
9.
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
10.
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
11.
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
12.
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
13.
Quartz and rutile were synthesized from silica-saturated aqueous fluids between 5 and 20 kbar and from 700 to 940°C in a piston-cylinder
apparatus to explore the potential pressure effect on Ti solubility in quartz. A systematic decrease in Ti-in-quartz solubility
occurs between 5 and 20 kbar. Titanium K-edge X-ray absorption near-edge structure (XANES) measurements demonstrate that Ti 4+ substitutes for Si 4+ on fourfold tetrahedral sites in quartz at all conditions studied. Molecular dynamic simulations support XANES measurements
and demonstrate that Ti incorporation onto fourfold sites is favored over interstitial solubility mechanisms. To account for
the P– T dependence of Ti-in-quartz solubility, a least-squares method was used to fit Ti concentrations in quartz from all experiments
to the simple expression
|
RTlnX\textTiO 2 \textquartz = - 60952 + 1.520 ·T(K) - 1741 ·P(kbar) + RTlna\textTiO 2 RT\ln X_{{{\text{TiO}}_{ 2} }}^{\text{quartz}} = - 60952 + 1.520 \cdot T(K) - 1741 \cdot P(kbar) + RT\ln a_{{{\text{TiO}}_{ 2} }} 相似文献
14.
We performed multi-anvil experiments in the system MgO-SiO 2 ± H 2O at 13.0–13.7 GPa and 1,025–1,300°C and in the system MgO-FeO-SiO 2 ± H 2O, under reducing conditions, at 11.0–12.7 GPa and 1,200°C, to depict the effect of H 2O 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-SiO 2-H 2O 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-SiO 2-H 2O 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. Fe 3+ concentrations as determined by EELS are very low and are expected to play no role in the broadening of the loops. 相似文献
15.
Bottom-water hypoxia effects on sediment–water interface nitrogen (N) transformations in Corpus Christi Bay (TX, USA) were
examined using continuous-flow intact sediment core incubations. Sediment cores were collected from three sites in August
2002 (summer hypoxia) and April 2003 (normoxia). Oxygen (O 2) and hydrogen sulfide (H 2S) depth profiles were generated with microelectrodes. Membrane inlet mass spectrometry was used to measure sediment O 2 demand and net N 2 flux and combined with isotope pairing to determine potential denitrification and N fixation. Potential dissimilatory nitrate
reduction to ammonium (DNRA) was measured using high-performance liquid chromatography. Sediment O 2 penetration depths ranged from 5 to 10 mm. H 2S ranged from being present in overlying water and throughout the sediment column in August to not detectable in overlying
water or sediment in April. Sediment O 2 demand was higher during bottom-water normoxia conditions versus hypoxia. Sediments were a significant source of
\text NH\text4\text + {\text{NH}}_{\text{4}}^{\text{ + }} to overlying water during hypoxia but not during normoxia. Net N 2 fixation was observed at one station in August and all stations in April. Denitrification rates were significantly higher
during hypoxia at two of three sites. Potential DNRA was observed during both oxic states, but rates were significantly higher
during hypoxia, which may reflect sulfide enhancement and absence of cation exchange with
\text14 \text NH\text4\text + ^{{\text{14}}} {\text{NH}}_{\text{4}}^{\text{ + }} . DNRA may contribute to formation and maintenance of bottom-water hypoxic events in this system. These results show that
N transformation pathways and rates change when bottom-water O 2 concentrations drop to hypoxic levels. Since south Texas is a semiarid region with few episodic runoff events, these results
indicate that Corpus Christi Bay sediments are a N source most of the year, and denitrification may drive N limitation between
episodic runoff events. 相似文献
16.
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
17.
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
18.
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. 相似文献
19.
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
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