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1.
A. Bhattacharya L. Mohanty A. Maji S. K. Sen M. Raith 《Contributions to Mineralogy and Petrology》1992,111(1):87-93
The existing experimental data [Ferry and Spear 1978; Perchuk and Lavrent'eva 1983] on Mg?Fe partitioning between garnet and biotite are disparate. The underlying assumption of ideal Mg?Fe exchange between the minerals has been examined on the basis of recently available thermochemical data. Using the updated mixing parameters for the pyrope-almandine asymmetric regular solution as inputs [Ganguly and Saxena 1984; Hackler and Wood 1984], thermodynamic analysis points to non-ideal mixing in the phlogopite-annite binary in the temperature range of 550°C–950°C. The non-ideality can be approximated by a temperature-independent, one constant Margules parameter. The retrieved values for enthalpy of mixing for Mg?Fe biotites and the standard state enthalpy and entropy changes of the exchange reaction were combined with existing thermochemical data on grossular-pyrope and grossular-almandine binaries to obtain geothermometric expressions for Mg?Fe fractionation between biotite and garnet. [T in K] $$\begin{gathered} {\text{T(HW) = [20286 + 0}}{\text{.0193P - \{ 2080(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{)}}^{\text{2}} {\text{ - 6350(X}}_{{\text{Fe}}}^{{\text{Gt}}} {\text{)}}^{\text{2}} \hfill \\ {\text{ - 13807(X}}_{{\text{Ca}}}^{{\text{Gt}}} {\text{)(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} {\text{) + 8540(X}}_{{\text{Fe}}}^{{\text{Gt}}} {\text{)(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{)(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} {\text{)}} \hfill \\ {\text{ + 4215(X}}_{{\text{Ca}}}^{{\text{Gt}}} {\text{)(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{ - X}}_{{\text{Fe}}}^{{\text{Gt}}} {\text{)\} + 4441}}{{{\text{(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} \mathord{\left/ {\vphantom {{{\text{(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} {{\text{[13}}{\text{.138}}}}} \right. \kern-\nulldelimiterspace} {{\text{[13}}{\text{.138}}}} \hfill \\ {\text{ + 8}}{\text{.3143 InK}}_{\text{D}} {\text{ + 6}}{\text{.276(X}}_{{\text{Ca}}}^{{\text{Gt}}} ){\text{(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} )] \hfill \\ {\text{T(GS) = [13538 + 0}}{\text{.0193P - \{ 837(X}}_{{\text{Mg}}}^{{\text{Gt}}} )^{\text{2}} {\text{ - 10460(X}}_{{\text{Fe}}}^{{\text{Gt}}} )^2 \hfill \\ {\text{ - 13807(X}}_{{\text{Ca}}}^{{\text{Gt}}} )(1{\text{ - X}}_{{\text{Mn}}}^{{\text{Gt}}} {\text{) + 19246(X}}_{{\text{Fe}}}^{{\text{Gt}}} ){\text{(X}}_{{\text{Mg}}}^{{\text{Gt}}} ){\text{(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} ) \hfill \\ {\text{ }}{{{\text{ + 5649(X}}_{{\text{Ca}}}^{{\text{Gt}}} ){\text{(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{ - X}}_{{\text{Fe}}}^{{\text{Gt}}} ){\text{\} + 7972(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} \mathord{\left/ {\vphantom {{{\text{ + 5649(X}}_{{\text{Ca}}}^{{\text{Gt}}} ){\text{(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{ - X}}_{{\text{Fe}}}^{{\text{Gt}}} ){\text{\} + 7972(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} {{\text{[6}}{\text{.778}}}}} \right. \kern-\nulldelimiterspace} {{\text{[6}}{\text{.778}}}} \hfill \\ {\text{ + 8}}{\text{.3143InK}}_{\text{D}} {\text{ + 6}}{\text{.276(X}}_{{\text{Ca}}}^{{\text{Gt}}} )(1{\text{ - X}}_{{\text{Mn}}}^{{\text{Gt}}} )] \hfill \\ \end{gathered} $$ The reformulated geothermometer is an improvement over existing biotite-garnet geothermometers because it reconciles the experimental data sets on Fe?Mg partitioning between the two phases and is based on updated activity-composition relationship in Fe?Mg?Ca garnet solid solutions. 相似文献
2.
S. K. Saxena 《Contributions to Mineralogy and Petrology》1979,70(3):229-235
A garnet-clinopyroxene geothermometer based on the available experimental data on compositions of coexisting phases in the system MgO-FeO-MnO-Al2O3-Na2O-SiO2 is as follows: $$T({\text{}}K) = \frac{{8288 + 0.0276 P {\text{(bar)}} + Q1 - Q2}}{{1.987 \ln K_{\text{D}} + 2.4083}}$$ where P is pressure, and Q1, Q2, and K D are given by the following equations $$Q1 = 2,710{\text{(}}X_{{\text{Fe}}} - X_{{\text{Mg}}} {\text{)}} + 3,150{\text{ }}X_{{\text{Ca}}} + 2,600{\text{ }}X_{{\text{Mn}}} $$ (mole fractions in garnet) $$\begin{gathered}Q2 = - 6,594[X_{{\text{Fe}}} {\text{(}}X_{{\text{Fe}}} - 2X_{{\text{Mg}}} {\text{)]}} \hfill \\{\text{ }} - 12762{\text{ [}}X_{{\text{Fe}}} - X_{{\text{Mg}}} (1 - X_{{\text{Fe}}} {\text{)]}} \hfill \\{\text{ }} - 11,281[X_{{\text{Ca}}} (1 - X_{{\text{Al}}} ) - 2X_{{\text{Mg}}} 2X_{{\text{Ca}}} ] \hfill \\{\text{ + 6137[}}X_{{\text{Ca}}} (2X_{{\text{Mg}}} + X_{{\text{Al}}} )] \hfill \\{\text{ + 35,791[}}X_{{\text{Al}}} (1 - 2X_{{\text{Mg}}} )] \hfill \\{\text{ + 25,409[(}}X_{{\text{Ca}}} )^2 ] - 55,137[X_{{\text{Ca}}} (X_{{\text{Mg}}} - X_{{\text{Fe}}} )] \hfill \\{\text{ }} - 11,338[X_{{\text{Al}}} (X_{{\text{Fe}}} - X_{{\text{Mg}}} )] \hfill \\\end{gathered} $$ [mole fractions in clinopyroxene Mg = MgSiO3, Fe = FeSiO3, Ca = CaSiO3, Al = (Al2O3-Na2O)] K D = (Fe/Mg) in garnet/(Fe/Mg) in clinopyroxene. Mn and Cr in clinopyroxene, when present in small concentrations are added to Fe and Al respectively. Fe is total Fe2++Fe3+. 相似文献
3.
Tjerk Peters Hans Schwander Volkmar Trommsdorff 《Contributions to Mineralogy and Petrology》1973,42(4):325-332
Reactions involving the phases quartz-rhodochrosite-tephroite-pyroxmangite-fluid have been studied experimentally in the system MnO-SiO2-CO2-H2O at a pressure of 2 000 bars and resulted in the following expressions 1 $$\begin{gathered} {\text{Rhodochrosite + Quartz = Pyroxmangite + CO}}_2 \hfill \\ {\text{ log}}_{{\text{10}}} K^{{\text{2000 bars}}} = - \frac{{11.765}}{T} + 18.618. \hfill \\ {\text{Rhodochrosite + Pyroxmangite = Tephroite + CO}}_2 \hfill \\ {\text{ log}}_{{\text{10}}} K^{{\text{2000 bars}}} = - \frac{{7.083}}{T} + 11.870. \hfill \\ \end{gathered}$$ which can be used to derive data for the remaining two reactions among the phases under consideration. Field data from the Alps are in agreement with the metamorphic sequence resulting from the experiments. 相似文献
4.
P. Bertrand C. Sotin J.-C. C. Mercier E. Takahashi 《Contributions to Mineralogy and Petrology》1986,93(2):168-178
The available experimental data on garnet-bearing-assemblages for synthetic chemical systems (MAS, FMAS, CMAS) have been used to calibrate consistent models for the Al-solubility in orthopyroxene coexisting with garnet, on the basis of equilibrium reaction Py(opx) ? Py(gt). The alternative reaction En(opx)+MgTs(opx) ? Py(gt) is discarded as it yields larger a-posteriori uncertainties. To provide a reliable equation, directly applicable to natural garnet lherzolites, each successive synthetic-system calibration is tested against Mori and Green's (1978) natural-system reequilibration data. For the MAS system, an ideal solution model with constant ΔH°, ΔV° and ΔS° based on 12-oxygen structural formulae for aluminous pyroxenes yields the best fit (GPa, K), $${\text{25,134 + 9,941 }}P - 23.177{\text{ }}T{\text{ + }}RT{\text{ ln (}}X_{{\text{Al}}}^{TB'} {\text{) = 0}}$$ . The MAS synthetic-system calibration can be directly applied to the FMAS system by adding an empirical correction term (20,835 [X Fe gt ]2) independent of either pressure and temperature. However, this correction term is not important because of the limited Fe content of mantle peridotites. When calcium is added to the MAS system, the equilibrium constant is calculated as: $$K_{{\text{CMAS}}} = {{[(1 - X_{{\text{Ca}}}^{M2} )^2 (X_{{\text{Al}}}^{TB'} )]} \mathord{\left/ {\vphantom {{[(1 - X_{{\text{Ca}}}^{M2} )^2 (X_{{\text{Al}}}^{TB'} )]} {[(1 - X_{{\text{Ca}}}^X )^3 (X_{{\text{Al}}}^Y )^2 ]}}} \right. \kern-\nulldelimiterspace} {[(1 - X_{{\text{Ca}}}^X )^3 (X_{{\text{Al}}}^Y )^2 ]}}$$ where M2 and TB′ are pyroxene sites and X and Y are garnet sites. Up to 5 GPa, X Ca X ～ and the CMAS experimental data agree well with the MAS model, but for Yamada and Takahashi's (1983) higher pressure experiments (up to 10 GPa), this no longer holds. Indeed, the garnet solid solution does not behave ideally and an asymmetric regular solution model is needed for application to the deepest natural samples available (>7GPa). Calibration based on new high pressure data yields, $$\begin{gathered} \Delta G_{{\text{CMAS}}}^{XS} = (X_{{\text{Ca}}}^X )(1 - X_{{\text{Ca}}}^X )(0.147 - X_{{\text{Ca}}}^X ) \hfill \\ {\text{ }} \cdot {\text{(6,440,535 - 1,490,654 }}P{\text{)}} \hfill \\ \end{gathered}$$ . According to tests of the inferred solution model, the CFMAS system is a good analogue of natural systems in the pressure, temperature and composition ranges covered by the natural-system reequilibration data (up to 1,500° C and 4 GPa). Simultaneous application of this thermobarometer and of the two-pyroxene mutual solubility thermometer (Bertrand and Mercier 1985) to the phases of the garnet-peridotite xenoliths from Thaba Putsoa, Lesotho, yields a refined paleogeotherm for southern Africa strongly contrasting with previous results. The “granular” nodules yield a thermal gradient of about 8 K/km characteristic of a lithospheric-type environment, whereas the “sheared” ones show a lower gradient of about 1 K/km. This is a typical geotherm expected for a steady thermal state with an inflexion point at the depth of about 160 km corresponding to the lithosphere/asthenosphere boundary. 相似文献
5.
The enstatite-diopside solvus presents certain interesting thermodynamic and crystal-structural problems. The solvus may be considered as parts of two solvi one with the ortho-structure and the other with clino-structure. By assuming the standard free energy change for the two reactions (MgMgSi2O6)opx ? (MgMgSi2O6)cpx and (CaMgSi2O6) opx ? (CaMgSi2O6) cpx as 500 and 1 000 to 3 000 cal/mol respectively, it is possible to calculate the regular solution parameter W for orthopyroxene and clinopyroxene. These W's essentially refer to mixing on M2 sites. The expression for the equilibrium constant by assuming ideal mixing for Fe-Mg, Fe-Ca and non-ideal mixing for Ca-Mg on binary M1 and ternary M2 sites is given by 1 $$K_a = \frac{{X_{{\text{Mg - cpx}}}^{{\text{M1}}} X_{{\text{Mg - cpx}}}^{{\text{M2}}} \exp \left[ {\frac{{W_{{\text{cpx}}} }}{{RT}}\left\{ {X_{{\text{Ca - cpx}}}^{{\text{M2}}} \left( {X_{{\text{Ca - cpx}}}^{{\text{M2}}} + X_{{\text{Fe - cpx}}}^{{\text{M2}}} } \right)} \right\}} \right]}}{{X_{{\text{Mg - cpx}}}^{{\text{M1}}} X_{{\text{Mg - opx}}}^{{\text{M2}}} \exp \left[ {\frac{{W_{{\text{cpx}}} }}{{RT}}\left\{ {X_{{\text{Ca - opx}}}^{{\text{M2}}} \left( {X_{{\text{Ca - opx}}}^{{\text{M2}}} + X_{{\text{Fe - opx}}}^{{\text{M2}}} } \right)} \right\}} \right]}}$$ where X's are site occupancies, R is 1.987 and T is temperature in oK. Temperature of pyroxene crystallization may be estimated by substituting for T in the above equation until the equation ?RT In K a=500 is satisfied. The shortcomings of this method are the incomplete standard free energy data on the end member components and the absence of site occupancy data in pyroxenes at high temperatures. The assumed free energy data do, however, show the possible extent of inaccuracy in temperature estimates resulting from the neglect of Mg-Ca non ideality. 相似文献
6.
Equilibrium alumina contents of orthopyroxene coexisting with spinel and forsterite in the system MgO-Al2O3-SiO2 have been reversed at 15 different P-T conditions, in the range 1,030–1,600° C and 10–28 kbar. The present data and three reversals of Danckwerth and Newton (1978) have been modeled assuming an ideal pyroxene solid solution with components Mg2Si2O6 (En) and MgAl2SiO6 (MgTs), to yield the following equilibrium condition (J, bar, K): $$\begin{gathered} RT{\text{ln(}}X_{{\text{MgTs}}} {\text{/}}X_{{\text{En}}} {\text{) + 29,190}} - {\text{13}}{\text{.42 }}T + 0.18{\text{ }}T + 0.18{\text{ }}T^{1.5} \hfill \\ + \int\limits_1^P {\Delta V_{T,P}^{\text{0}} dP = 0,} \hfill \\ \end{gathered} $$ where $$\begin{gathered} + \int\limits_1^P {\Delta V_{T,P}^{\text{0}} dP} \hfill \\ = [0.013 + 3.34 \times 10^{ - 5} (T - 298) - 6.6 \times 10^{ - 7} P]P. \hfill \\ \end{gathered} $$ The data of Perkins et al. (1981) for the equilibrium of orthopyroxene with pyrope have been similarly fitted with the result: $$\begin{gathered} - RT{\text{ln(}}X_{{\text{MgTs}}} \cdot X_{{\text{En}}} {\text{) + 5,510}} - 88.91{\text{ }}T + 19{\text{ }}T^{1.2} \hfill \\ + \int\limits_1^P {\Delta V_{T,P}^{\text{0}} dP = 0,} \hfill \\ \end{gathered} $$ where $$\begin{gathered} + \int\limits_1^P {\Delta V_{T,P}^{\text{0}} dP} \hfill \\ = [ - 0.832 - 8.78{\text{ }} \times {\text{ 10}}^{ - {\text{5}}} (T - 298) + 16.6{\text{ }} \times {\text{ 10}}^{ - 7} P]{\text{ }}P. \hfill \\ \end{gathered} $$ The new parameters are in excellent agreement with measured thermochemical data and give the following properties of the Mg-Tschermak endmember: $$H_{f,970}^0 = - 4.77{\text{ kJ/mol, }}S_{298}^0 = 129.44{\text{ J/mol}} \cdot {\text{K,}}$$ and $$V_{298,1}^0 = 58.88{\text{ cm}}^{\text{3}} .$$ The assemblage orthopyroxene+spinel+olivine can be used as a geothermometer for spinel lherzolites, subject to a choice of thermodynamic mixing models for multicomponent orthopyroxene and spinel. An ideal two-site mixing model for pyroxene and Sack's (1982) expressions for spinel activities provide, with the present experimental calibration, a geothermometer which yields temperatures of 800° C to 1,350° C for various alpine peridotites and 850° C to 1,130° C for various volcanic inclusions of upper mantle origin. 相似文献
7.
The Gibbs free energy and volume changes attendant upon hydration of cordierites in the system magnesian cordierite-water have been extracted from the published high pressure experimental data at \(P_{{\text{H}}_{\text{2}} {\text{O}}} \) =P total, assuming an ideal one site model for H2O in cordierite. Incorporating the dependence of ΔG and ΔV on temperature, which was found to be linear within the experimental conditions of 500°–1,000°C and 1–10,000 bars, the relation between the water content of cordierite and P, T and \(f_{{\text{H}}_{\text{2}} {\text{O}}} \) has been formulated as $$\begin{gathered} X_{{\text{H}}_{\text{2}} {\text{O}}}^{{\text{crd}}} = \hfill \\ \frac{{f_{{\text{H}}_{\text{2}} {\text{O}}}^{{\text{P, T}}} }}{{\left[ {{\text{exp}}\frac{1}{{RT}}\left\{ {64,775 - 32.26T + G_{{\text{H}}_{\text{2}} {\text{O}}}^{{\text{1, }}T} - P\left( {9 \times 10^{ - 4} T - 0.5142} \right)} \right\}} \right] + f_{{\text{H}}_{\text{2}} {\text{O}}}^{{\text{P, T}}} }} \hfill \\ \end{gathered} $$ The equation can be used to compute H2O in cordierites at \(P_{{\text{H}}_{\text{2}} {\text{O}}} \) <1. Our results at different P, T and partial pressure of water, assuming ideal mixing of H2O and CO2 in the vapour phase, are in very good agreement with the experimental data of Johannes and Schreyer (1977, 1981). Applying the formulation to determine \(X_{{\text{H}}_{\text{2}} {\text{O}}}^{{\text{crd}}} \) in the garnet-cordierite-sillimanite-plagioclase-quartz granulites of Finnish Lapland as a test case, good agreement with the gravimetrically determined water contents of cordierite was obtained. Pressure estimates, from a thermodynamic modelling of the Fe-cordierite — almandine — sillimanite — quartz equilibrium at \(P_{{\text{H}}_{\text{2}} {\text{O}}} = 0\) and \(P_{{\text{H}}_{\text{2}} {\text{O}}} \) =Ptotal, for assemblages from South India, Scottish Caledonides, Daly Bay and Hara Lake areas are compatible with those derived from the garnetplagioclase-sillimanite-quartz geobarometer. 相似文献
8.
The effective binary diffusion coefficient (EBDC) of silicon has been measured during the interdiffusion of peralkaline, fluorine-bearing (1.3 wt% F), hydrous (3.3 and 6 wt% H2O), dacitic and rhyolitic melts at 1.0 GPa and temperatures between 1100°C and 1400°C. From Boltzmann-Matano analysis of diffusion profiles the diffusivity of silicon at 68 wt% SiO2 can be described by the following Arrhenius equations (with standard errors): $$\begin{gathered} {\text{with 1}}{\text{.3 wt\% F and 3}}{\text{.3\% H}}_{\text{2}} {\text{O:}} \hfill \\ {\text{D}}_{{\text{Si}}} = \begin{array}{*{20}c} { + {\text{3}}{\text{.59}}} \\ {{\text{3}}{\text{.66}} \times {\text{10}}^{ - {\text{9}}} } \\ { - {\text{1}}{\text{.86}}} \\ \end{array} {\text{exp}}\left( {{{ - {\text{86}}{\text{.1}} \pm {\text{8}}{\text{.9}}} \mathord{\left/ {\vphantom {{ - {\text{86}}{\text{.1}} \pm {\text{8}}{\text{.9}}} {{\text{RT}}}}} \right. \kern-\nulldelimiterspace} {{\text{RT}}}}} \right) \hfill \\ {\text{with 1}}{\text{.3 wt\% F and 6}}{\text{.0\% H}}_{\text{2}} {\text{O:}} \hfill \\ {\text{D}}_{{\text{Si}}} = \begin{array}{*{20}c} { + {\text{3}}{\text{.59}}} \\ {{\text{3}}{\text{.51}} \times {\text{10}}^{ - {\text{8}}} } \\ { - {\text{1}}{\text{.77}}} \\ \end{array} {\text{exp}}\left( {{{ - {\text{109}}{\text{.5}} \pm {\text{8}}{\text{.9}}} \mathord{\left/ {\vphantom {{ - {\text{109}}{\text{.5}} \pm {\text{8}}{\text{.9}}} {{\text{RT}}}}} \right. \kern-\nulldelimiterspace} {{\text{RT}}}}} \right) \hfill \\ \end{gathered} $$ where D is in m2s?1 and activation energies are in kJ/mol. Diffusivities measured at 64 and 72 wt% SiO2 are only slightly different from those at 68 wt% SiO2 and frequently all measurements are within error of each other. Silicon, aluminum, iron, magnesium, and calcium EBDCs were also calculated from diffusion profiles by error function inversion techniques assuming constant diffusivity. With one exception, silicon EBDCs calculated by error function techniques are within error of Boltzmann-Matano EBDCs. Average diffusivities of Fe, Mg, and Ca were within a factor of 2.5 of silicon diffusivities whereas Al diffusivities were approximately half those of silicon. Alkalies diffused much more rapidly than silicon and non-alkalies, however their diffusivities were not quantitatively determined. Low activation energies for silicon EBDCs result in rapid diffusion at magmatic temperatures. Assuming that water and fluorine exert similar effects on melt viscosity at high temperatures, the viscosity can be calculated and used in the Eyring equation used to determine diffusivities, typically to within a factor of three of those measured in this study. This correlation between viscosity and diffusivity can be inverted to calculate viscosities of fluorine- and water-bearing granitic melts at magmatic temperatures; these viscosities are orders of magnitude below those of hydrous granitic melts and result in more rapid and effective separation of granitic magmas from partially molten source rocks. Comparison of Arrhenius parameters for diffusion measured in this study with Arrhenius parameters determined for diffusion in similar compositions at the same pressure demonstrates simple relationships between Arrhenius parameters, activation energy-Ea, kJ/mol, pre-exponential factor-Do, m2s?1, and the volatile, X=F or OH?, to oxygen, O, ratio of the melt {(X/X+O)}: $$\begin{gathered} {\text{E}}a = - {\text{1533\{ }}{{\text{X}} \mathord{\left/ {\vphantom {{\text{X}} {\left( {{\text{X}} + {\text{O}}} \right)}}} \right. \kern-\nulldelimiterspace} {\left( {{\text{X}} + {\text{O}}} \right)}}{\text{\} }} + {\text{213}}{\text{.3}} \hfill \\ {\text{D}}_{\text{O}} = {\text{2}}{\text{.13}} \times {\text{10}}^{ - {\text{6}}} {\text{exp}}\left[ { - {\text{6}}{\text{.5\{ }}{{\text{X}} \mathord{\left/ {\vphantom {{\text{X}} {\left( {{\text{X}} + {\text{O}}} \right)}}} \right. \kern-\nulldelimiterspace} {\left( {{\text{X}} + {\text{O}}} \right)}}{\text{\} }}} \right] \hfill \\ \end{gathered} $$ These relationships can be used to estimate diffusion in various melts of dacitic to rhyolitic composition containing both fluorine and water. Calculations for the contamination of rhyolitic melts by dacitic enclaves at 800°C and 700°C provide evidence for the virtual inevitability of diffusive contamination in hydrous and fluorine-bearing magmas if they undergo magma mixing of any form. 相似文献
9.
Partitioning of Mg and Fe between coexisting biotite and orthopyroxene has been experimentally determined at temperatures 700, 750 and 800° C and 490 MPa total pressure in the system KAlO2-MgO-FeO-SiO2-H2O. Oxygen fugacity was controlled by the QFM buffer. Starting materials were synthetic minerals of differing Fe/(Fe+Mg) values. Run products were analyzed for partitioning of components by a microprobe. Orthopyroxene was established to be notably inhomogeneous, whereas biotite was essentially homogeneous. To establish equilibrium relations, statistical treatment of the results of each experiment in addition to the whole complex of experimental data was applied. The regression equations for isotherms of the Fe-Mg partitioning between the minerals studied have been obtained. As a result, the equation for a two-dimensional regression may be written as: $$\begin{gathered} Y = (A + A_1 t + A_2 t^2 )(X - X^4 ) + (B + B_1 t + B_1 t^2 )(X^2 - X^4 ) + \hfill \\ (C + C_1 t + C_1 t^2 )(X^3 - X^4 ) + X^4 {\text{ where }}Y = X_{{\text{Opx}}}^{{\text{Fe}}} ;{\text{ X}} = {\text{X}}_{{\text{Bi}}}^{{\text{Fe}}} ; \hfill \\ t = 1000/T,K, \hfill \\ \begin{array}{*{20}c} {A = {\text{ }}4.59398,} & {A_1 = - {\text{ }}8.29838,} & {A_2 = {\text{ }}4.97316,} \\ {B = - 11.13731,} & {B_1 = {\text{ }}28.19304,} & {B_2 = - 20.98240,} \\ {A = {\text{ }}8.25072,} & {C_1 = - 20.80485,} & {C_2 = {\text{ }}15.35967} \\ \end{array} \hfill \\ {\text{ }}\sigma = 0.0143{\text{ }} \hfill \\ \end{gathered}$$ . This equation enables extrapolation of partitioning isotherms over a wide range of temperatures. 相似文献
10.
A mineralogic geobarometer based on the reaction garnet+clinopyroxene+quartz=2 orthopyroxene+anorthite is proposed. The geobarometric formulations for the Fe- and Mg- end member equilibria are $$\begin{gathered} P_{({\text{Fe}})} {\text{ }}({\text{bars}}){\text{ = 32}}{\text{.097 }}T{\text{ }} - {\text{ 26385 }} - {\text{ 22}}{\text{.79 (}}T - 848 - T1{\text{n(}}T/848{\text{))}} \hfill \\ {\text{ }} - (3.655 + 0.0138T){\text{ }}\left( {\frac{{{\text{(}}T - 848{\text{)}}^{\text{2}} }}{T}} \right) \hfill \\ {\text{ }} - {\text{(3}}{\text{.123) }}T1{\text{n }}\frac{{(a_{a{\text{n}}}^{{\text{Plag}}} )(a_{{\text{fs}}}^{{\text{P}}\ddot u{\text{x}}} )^2 }}{{(a_{{\text{alm}}}^{{\text{Gt}}} )(a_{{\text{hed}}}^{{\text{Opx}}} )}} \hfill \\ P_{({\text{Mg}})} {\text{ (bars) = 9}}{\text{.270 }}T + 4006 - 0.9305{\text{ }}(T - 848 - T1{\text{n (}}T/848{\text{)}}) \hfill \\ {\text{ }} - (1.1963{\text{ }} - {\text{ }}6.0128{\text{ x 10}}^{ - {\text{3}}} T)\left( {\frac{{(T - 848)^2 }}{T}} \right) \hfill \\ {\text{ }} - 3.489{\text{ }}T1{\text{n }}\frac{{(a_{an}^{{\text{Plag}}} ){\text{ }}(a_{{\text{ens}}}^{{\text{Opx}}} )}}{{{\text{(}}a_{{\text{pyr}}}^{{\text{Gt}}} {\text{) (}}a_{{\text{diop}}}^{{\text{Cpx}}} {\text{)}}}}. \hfill \\ \end{gathered}$$ The end member thermodynamic data have been taken from the data base of Helgeson et al. (1978) and Saxena and Erikson (1983). The activities of pyroxene components and anorthite in plagioclase have been modelled after Wood and Banno (1973) and Newton (1983) respectively. The activities of pyrope and almandine are calculated from the binary interaction parameters for garnet solid solutions proposed by Saxena and Erikson (1983). Pressures computed from these equations for fifty sets of published mineral data from several granulite areas are comparable with those obtained from dependable geobarometers. The pressure values determined from the Fe-end member equilibrium appear to be more reasonable than those from the Mg-end member reaction. It is likely that the difference in pressures computed from the Fe- and Mg-end members, ΔP *, have been caused by non-ideal mixing in the phases, especially in garnets. 相似文献
11.
The enthalpy of formation of andradite (Ca3Fe2Si3O12) has been estimated as-5,769.700 (±5) kJ/mol from a consideration of the calorimetric data on entropy (316.4 J/mol K) and of the experimental phaseequilibrium data on the reactions: 1 $$\begin{gathered} 9/2 CaFeSi_2 O_6 + O_2 = 3/2 Ca_3 Fe_2 Si_3 O_{12} + 1/2 Fe_3 O_4 + 9/2 SiO_2 (a) \hfill \\ Hedenbergite andradite magnetite quartz \hfill \\ \end{gathered} $$ 1 $$\begin{gathered} 4 CaFeSi_2 O_6 + 2 CaSiO_3 + O_2 = 2 Ca_3 Fe_2 Si_3 O_{12} + 4 SiO_2 (b) \hfill \\ Hedenbergite wollastonite andradite quartz \hfill \\ \end{gathered} $$ 1 $$\begin{gathered} 18 CaSiO_3 + 4 Fe_3 O_4 + O_2 = 6Ca_3 Fe_2 Si_3 O_{12} (c) \hfill \\ Wollastonite magnetite andradite \hfill \\ \end{gathered} $$ 1 $$\begin{gathered} Ca_3 Fe_2 Si_3 O_{12} = 3 CaSiO_3 + Fe_2 O_3 . (d) \hfill \\ Andradite pseudowollastonite hematite \hfill \\ \end{gathered} $$ and $$log f_{O_2 } = E + A + B/T + D(P - 1)/T + C log f_{O_2 } .$$ Oxygen-barometric scales are presented as follows: $$\begin{gathered} E = 12.51; D = 0.078; \hfill \\ A = 3 log X_{Ad} - 4.5 log X_{Hd} ; C = 0; \hfill \\ B = - 27,576 - 1,007(1 - X_{Ad} )^2 - 1,476(1 - X_{Hd} )^2 . \hfill \\ \end{gathered} $$ For the assemblage andradite (Ad)-hedenbergite (Hd)-magnetite-quartz: $$\begin{gathered} E = 13.98; D = 0.0081; \hfill \\ A = 4 log(X_{Ad} / X_{Hd} ); C = 0; \hfill \\ B = - 29,161 - 1,342.8(1 - X_{Ad} )^2 - 1,312(1 - X_{Hd} )^2 . \hfill \\ \end{gathered} $$ For the assemblage andradite-hedenbergite-wollastonite-quartz: 1 $$\begin{gathered} E = 13.98;{\text{ }}D = 0.0081; \hfill \\ A = 4\log (X_{Ad} /X_{Hd} );{\text{ C = 0;}} \hfill \\ B = - 29,161 - 1,342.8(1 - X_{Ad} )^2 - 1,312(1 - X_{Hd} )^2 . \hfill \\ \end{gathered} $$ For the assemblage andradite-hedenbergite-calcitequartz: 1 $$\begin{gathered} E = - 1.69;{\text{ }}D = - 0.199; \hfill \\ A = 4\log (X_{Ad} /X_{Hd} );{\text{ C = 2;}} \hfill \\ B = - 20,441 - 1,342.8(1 - X_{Ad} )^2 - 1,312(1 - X_{Hd} )^2 . \hfill \\ \end{gathered} $$ For the assemblage andradite-hedenbergite-wollastonite-calcite: 1 $$\begin{gathered} E = - 17.36;{\text{ }}D = - 0.403; \hfill \\ A = 4\log (X_{Ad} /X_{Hd} );{\text{ C = 4;}} \hfill \\ B = - 11,720 - 1,342.8(1 - X_{Ad} )^2 - 1,312(1 - X_{Hd} )^2 \hfill \\ \end{gathered} $$ The oxygen fugacity of formation of those skarns where andradite and hedenbergite assemblage is typical can be calculated by using the above equations. The oxygen fugacity of formation of this kind of skarn ranges between carbon dioxide/graphite and hematite/magnetite buffers. It increases from the inside zones to the outside zones, and appears to decrease with the ore-types in the order Cu, Pb?Zn, Fe, Mo, W(Sn) ore deposits. 相似文献
12.
Yastami Oka Petra Steinke Niranjan D. Chatterjee 《Contributions to Mineralogy and Petrology》1984,87(2):196-204
Three Al-Cr exchange isotherms at 1,250°, 1,050°, and 796° between Mg(Al, Cr)2O4 spinel and (Al, Cr)2O3 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)2O3 crystalline solution (Chatterjee et al. 1982a) and assuming that the Mg(Al, Cr)2O4 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)2O4 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 dTc/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
13.
The system Fe-Si-O: Oxygen buffer calibrations to 1,500K 总被引：1，自引：0，他引：1
The five solid-phase oxygen buffers of the system Fe-Si-O, iron-wuestite (IW), wuestite-magnetite (WM), magnetite-hematite (MH), quartz-iron-fayalite (QIF) and fayalite-magnetite-quartz (FMQ) have been recalibrated at 1 atm pressure and temperatures from 800°–1,300° C, using a thermogravimetric gas mixing furnace. The oxygen fugacity, \(f_{{\text{O}}_{\text{2}} }\) was measured with a CaO-doped ZrO2 electrode. Measurements were made also for wuestite solid solutions in order to determine the redox behavior of wuestites with O/Fe ratios varying from 1.05 to 1.17. For FMQ, additional determinations were carried out at 1 kb over a temperature range of 600° to 800° C, using a modified Shaw membrane. Results agree reasonably well with published data and extrapolations. The reaction parameters K, ΔG r o , ΔH r o , and ΔS r o were calculated from the following log \(f_{{\text{O}}_{\text{2}} }\) /T relations (T in K): $$\begin{gathered} {\text{IW }}\log f_{{\text{O}}_{\text{2}} } = - 26,834.7/T + 6.471\left( { \pm 0.058} \right) \hfill \\ {\text{ }}\left( {{\text{800}} - 1,260{\text{ C}}} \right), \hfill \\ {\text{WM }}\log f_{{\text{O}}_{\text{2}} } = - 36,951.3/T + 16.092\left( { \pm 0.045} \right) \hfill \\ {\text{ }}\left( {{\text{1,000}} - 1,300{\text{ C}}} \right), \hfill \\ {\text{MH }}\log f_{{\text{O}}_{\text{2}} } = - 23,847.6/T + 13.480\left( { \pm 0.055} \right) \hfill \\ {\text{ }}\left( {{\text{1,040}} - 1,270{\text{ C}}} \right), \hfill \\ {\text{QIF }}\log f_{{\text{O}}_{\text{2}} } = - 27,517.5/T + 6.396\left( { \pm 0.049} \right) \hfill \\ {\text{ }}\left( {{\text{960}} - 1,140{\text{ C}}} \right), \hfill \\ {\text{FMQ }}\log f_{{\text{O}}_{\text{2}} } = - 24,441.9/T + 8.290\left( { \pm 0.167} \right) \hfill \\ {\text{ }}\left( {{\text{600}} - 1,140{\text{ C}}} \right). \hfill \\ \end{gathered}$$ These experimentally determined reaction parameters were combined with published 298 K data to determine the parameters Gf, Hf, and Sf for the phases wuestite, magnetite, hematite, and fayalite from 298 K to the temperatures of the experiments. The T? \(f_{{\text{O}}_{\text{2}} }\) data for wuestite solid solutions were used to obtain activities, excess free energies and Margules mixing parameters. The new data provide a more reliable, consistent and complete reference set for the interpretation of redox reactions at elevated temperatures in experiments and field settings encompassing the crust, mantle and core as well as extraterrestrial environments. 相似文献
14.
Solubility of Ca and Al in orthopyroxene from spinel peridotite: an improved version of an empirical geothermometer 总被引：1，自引：1，他引：0
Geothermometric equations for spinel peridotites by Fujii (1976), Gasparik and Newton (1984), and Chatterjee, and Terhart (1985) based on the reaction enstatite (en)+spinel (sp)Mg–Tschermaks (mats)+forsterite (fo) were tested using a nearly isothermal suite of mantle xenoliths from the Eifel, West Germany. In spite of using activities of MgAl2O4, en, and mats to allow for the non-ideal solution behaviour of the constituent phases, temperatures calculated from these equations systematically change as a function of Cr/(Cr+AL+Fe3+) in spinel. We propose an improved version of the empirical geothermometer for spinel peridotites of Sachtleben and Seck (1981) derived from the evaluation of the solubilities of Ca and Al in orthopyroxene from more than 100 spinel peridotites from the Rhenish Volcanic Province. A least squares regression yielded a smooth correlation between
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相似文献
15.
Mössbauer and polarized optical absorption spectra of the kyanite-related mineral yoderite were recorded. Mössbauer spectra of the purple (PY) and green yoderite (GY) from Mautia Hill, Tanzania, show that the bulk of the iron is Fe3+ in both varieties, with Fe2+/(Fe2++Fe3+) ratios near 0.05. Combining this result with new microprobe data for PY and with literature data for GY gives the crystallochemical formulae: $$\begin{gathered} ({\text{Mg}}_{{\text{1}}{\text{.95}}} {\text{Fe}}_{{\text{0}}{\text{.02}}}^{{\text{2 + }}} {\text{Mn}}_{{\text{0}}{\text{.01}}}^{{\text{2 + }}} {\text{Fe}}_{{\text{0}}{\text{.34}}}^{{\text{3 + }}} {\text{Mn}}_{{\text{0}}{\text{.07}}}^{{\text{3 + }}} {\text{Ti}}_{{\text{0}}{\text{.01}}} {\text{Al}}_{{\text{3}}{\text{.57}}} )_{5.97}^{[5,6]} \hfill \\ {\text{Al}}_{{\text{2}}{\text{.00}}}^{{\text{[5]}}} [({\text{Si}}_{{\text{3}}{\text{.98}}} {\text{P}}_{{\text{0}}{\text{.03}}} ){\text{O}}_{{\text{18}}{\text{.02}}} ({\text{OH)}}_{{\text{1}}{\text{.98}}} ] \hfill \\ \end{gathered}$$ and PY and $$\begin{gathered} ({\text{Mg}}_{{\text{1}}{\text{.98}}} {\text{Fe}}_{{\text{0}}{\text{.02}}}^{{\text{2 + }}} {\text{Mn}}_{{\text{< 0}}{\text{.001}}}^{{\text{2 + }}} {\text{Fe}}_{{\text{0}}{\text{.45}}}^{{\text{3 + }}} {\text{Ti}}_{{\text{0}}{\text{.01}}} {\text{Al}}_{{\text{3}}{\text{.56}}} )_{6.02}^{[5,6]} \hfill \\ {\text{Al}}_{{\text{2}}{\text{.00}}}^{{\text{[5]}}} [({\text{Si}}_{{\text{3}}{\text{.91}}} {\text{O}}_{{\text{17}}{\text{.73}}} {\text{(OH)}}_{{\text{2}}{\text{.27}}} ] \hfill \\ \end{gathered}$$ for GY. The Mössbauer spectra at room temperature contain one main doublet with isomer shifts and quadrupole splittings of 0.36 (PY), 0.38 (GY) and 1.00 (PY), 0.92 (GY) mm s?1, respectively. These values correspond to Fe3+ in six or five-fold coordination. The doublet components have anomalously large half widths indicating either accomodation of Fe3+ in more than one position (e.g., octahedraA1 and five coordinatedA2) or the yet unresolved superstructure. Besides strong absorption in the ultraviolet (UV) starting from about 25,000 cm?1, the polarized optical absorption spectra are dominated by strong bands around 16,500 and 21,000 cm?1 (PY) and a medium strong band at around 13,800 cm?1 (GY). Position and polarization of these bands, in combination with the UV absorption, explain the colour and pleochroism of the two varieties. The bands in question are assigned to homonuclear metal-to-metal charge transfer transitions: Mn2+(A1) Mn3+(A1′) ? Mn3+(A1) Mn2+(A1′) and Mn2+(A1) Mn3+(A2 ? Mn3+(A1) Mn2+(A2) in PY and Fe2+(A1) Fe3+(A1′) ? Fe3+(A1) Fe2+(A1′) in GY. The evidence for homonuclear Mn2+ Mn3+ charge transfer (CTF) is not quite clear and needs further study. Heteronuclear FeTi CTF does not contribute to the spectra. In PY, additional weak bands were resolved at energies around 17,700, 18,700, 21,000, and 21,900 cm?1 and assigned to Mn3+ in two positions. Weak bands around 10,000 cm?1 in both varieties are assigned to Fe2+ spin-alloweddd-transitions. Very weak and sharp bands, around 15,400, 16,400, 21,300, 22,100, 23,800, and 25,000 cm?1 are identified in GY and assigned to Fe3+ spin-forbiddendd-transitions. 相似文献
16.
Ephesite, Na(LiAl2) [Al2Si2O10] (OH)2, has been synthesized for the first time by hydrothermal treatment of a gel of requisite composition at 300≦T(° C)≦700 and \(P_{H_2 O}\) upto 35 kbar. At \(P_{H_2 O}\) between 7 and 35 kbar and above 500° C, only the 2M1 polytype is obtained. At lower temperatures and pressures, the 1M polytype crystallizes first, which then inverts to the 2M1 polytype with increasing run duration. The X-ray diffraction patterns of the 1M and 2M1 poly types can be indexed unambiguously on the basis of the space groups C2 and Cc, respectively. At its upper thermal stability limit, 2M1 ephesite decomposes according to the reaction
17.
The mixing properties of Fe3Al2Si3O12-Ca3Al2Si3O12 garnet solid solutions have been studied in the temperature range 850–1100° C. The experimental method involves measuring the composition of garnet in equilibrium with an assemblage in which the activity of the Ca3Al2Si3O12 component is fixed. Experiments on the assemblage garnet solid solution, anorthite, Al2SiO5 polymorph and quartz at known pressure and temperature fix the activity of the Ca3Al2Si3O12 component through the equilibrium:
18.
A. Bhattacharya 《Contributions to Mineralogy and Petrology》1986,94(3):387-394
Five geobarometers involving cordierite have been formulated for quantitative pressure sensing in high grade metapelites. The relevant reactions in the FeO-Al2O3-SiO2 (±H2O) system are based on the assemblages (A) cordierite-garnet-sillimanite-quartz, (B) cordierite-spinel-quartz, (C) cordierite-garnet-spinel-sillimanite, (D) cordierite-garnet-orthopyroxene-quartz and (E) cordierite-orthopyroxene-sillimanite-quartz. Application of the barometric formulations to a large number of granulite grade rocks indicates that the cordierite-garnet-sillimanite-quartz equilibrium is widely applicable and registers pressures which are in good agreement with the “consensus” pressure estimates. The dispersion in the computed P values, expressed as one standard deviation, is within ±1.2 kbar. The geobarometers (B) and (C) also yield pressures which are reasonable and compare well with those computed from equilibrium (A). The estimated pressures from (D) and (E), both involving orthopyroxene, are at variance with these estimates. It has been argued that the discrepancy in pressures obtained from these geobarometers stems from an inadequate knowledge of activity-composition relations and/or errors in input thermodynamic data of aluminous orthopyroxene. The convergence of pressure values estimated from the barometric formulations, especially (A), (B) and (C), implies that the present formulations are more dependable than the existing formulations and are also capable of setting limits on P values in response to varying $$\begin{gathered} {\text{1/2Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} \hfill \\ {\text{ = 1/3Fe}}_{\text{3}} {\text{Al}}_{\text{2}} {\text{Si}}_{\text{3}} {\text{O}}_{{\text{12}}} {\text{ + 2/3Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{ + 5/6SiO}}_{\text{2}} {\text{. (A)}} \hfill \\ {\text{1/2Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} {\text{ = FeAl}}_{\text{2}} {\text{O}}_{\text{4}} {\text{ + 5/2SiO}}_{\text{2}} {\text{. (B)}} \hfill \\ {\text{Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} {\text{ + FeAl}}_{\text{2}} {\text{O}}_{\text{4}} \hfill \\ = {\text{Fe}}_{\text{3}} {\text{Al}}_{\text{2}} {\text{Si}}_{\text{3}} {\text{O}}_{{\text{12}}} {\text{ + 2Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{. (C)}} \hfill \\ {\text{1/2Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} {\text{ + Fe}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} \hfill \\ = {\text{Fe}}_{\text{3}} {\text{Al}}_{\text{2}} {\text{Si}}_{\text{3}} {\text{O}}_{{\text{12}}} {\text{ + 3/2SiO}}_{\text{2}} .{\text{ (D)}} \hfill \\ {\text{1/2Fe}}_{\text{2}} {\text{Al}}{}_{\text{4}}{\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} \hfill \\ = 1/2{\text{Fe}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} {\text{ + Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{ + 1/2SiO}}_{\text{2}} .{\text{ (E)}} \hfill \\ \end{gathered}$$ . The present communication addresses the calibration, applicability and reliability of these barometers with reference to granulite facies metapelites. 相似文献
19.
Detailed analysis of textural and chemical criteria in rocks of the anorthosite-charnockite suite of the Adirondack Highlands suggests that development of garnet in silica-saturated rocks of the suite occurs according to the reaction: $$\begin{gathered} {\text{Anorthite}} {\text{Orthopyroxene}} {\text{Quartz}} \hfill \\ {\text{2CaAl}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{8}} + (6 - \alpha )({\text{Fe,Mg}}){\text{SiO}}_{\text{3}} + \alpha {\text{Fe - Oxide + (}}\alpha {\text{ - 2)SiO}}_{\text{2}} \hfill \\ {\text{Garnet}} {\text{Clinopyroxene}} \hfill \\ = {\text{Ca(Fe,Mg)}}_{\text{5}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{6}} {\text{O}}_{{\text{24}}} + {\text{Ca(Fe,Mg)Si}}_{\text{2}} {\text{O}}_{\text{6}} \hfill \\ \end{gathered} $$ , where α is a function of the distribution of Fe and Mg between the several coexisting ferromagnesian phases. Depending upon the relative amounts of Fe and Mg present, quartz may be either a reactant or a product. Using an aluminum-fixed reference frame, this reaction can be restated in terms of a set of balanced partial reactions describing the processes occurring in spatially separated domains within the rock. The fact that garnet invariably replaces plagioclase as opposed to the other reactant phases indicates that the aluminum-fixed model is valid as a first approximation. This reaction is univariant and produces unzoned garnet. It differs from a similar equation proposed by de Waard (1965) for the origin of garnet in Adirondack metabasic rocks, i.e. 6 Orthopyroxene+2 Anorthite = Clinopyroxene+Garnet+2 Quartz, the principle difference being that iron oxides (ilmenite and/or magnetite) are essential reactant phases in the present reactions. The product assemblage (garnet+clinopyroxene+plagioclase ± orthopyroxene ± quartz) is characteristic of the clinopyroxene-almandine subfacies of the granulite facies. 相似文献
20.
Tibor Gasparik 《Contributions to Mineralogy and Petrology》1985,89(4):346-357
In the system Na2O-CaO-Al2O3-SiO2 (NCAS), the equilibrium compositions of pyroxene coexisting with grossular and corundum were experimentally determined at 40 different P-T conditions (1,100–1,400° C and 20.5–38 kbar). Mixing properties of the Ca-Tschermak — Jadeite pyroxene inferred from the data are (J, K): $$\begin{gathered} G_{Px}^{xs} = X_{{\text{CaTs}}} X_{{\text{Jd}}} [14,810 - 7.15T - 5,070(X_{{\text{CaTs}}} - X_{{\text{Jd}}} ) \hfill \\ {\text{ }} - 3,350(X_{{\text{CaTs}}} - X_{{\text{Jd}}} )^2 ] \hfill \\ \end{gathered} $$ The excess entropy is consistent with a complete disorder of cations in the M2 and the T site. Compositions of coexisting pyroxene and plagioclase were obtained in 11 experiments at 1,190–1,300° C/25 kbar. The data were used to infer an entropy difference between low and high anorthite at 1,200° C, corresponding to the enthalpy difference of 9.6 kJ/mol associated with the C \(\bar 1\) =I \(\bar 1\) transition in anorthite as given by Carpenter and McConnell (1984). The resulting entropy difference of 5.0 J/ mol · K places the transition at 1,647° C. Plagioclase is modeled as ideal solutions, C \(\bar 1\) and I \(\bar 1\) , with a non-first order transition between them approximated by an empirical expression (J, bar, K): $$\Delta G_T = \Delta G_{1,473} \left[ {1 - 3X_{Ab} \tfrac{{T^4 - 1,473^4 }}{{\left( {1,920 - 0.004P} \right)^4 - 1,473^4 }}} \right],$$ where $$\Delta G_{1,473} = 9,600 - 5.0T - 0.02P$$ The derived mixing properties of the pyroxene and plagioclase solutions, combined with the thermodynamic properties of other phases, were used to calculate phase relations in the NCAS system. Equilibria involving pyroxene+plagioclase +grossular+corundum and pyroxene+plagioclase +grossular+kyani te are suitable for thermobarometry. Albite is the most stable plagioclase. 相似文献
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