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1.
The biotite zone assemblage: calcite-quartz-plagioclase (An25)-phengite-paragonite-chlorite-graphite, is developed at the contact between a carbonate and a pelite from British Columbia. Thermochemical data for the equilibrium paragonite+calcite+2 quartz=albite+ anorthite+CO2+H2O yields: $$\log f{\text{H}}_{\text{2}} {\text{O}} + \log f{\text{CO}}_{\text{2}} = 5.76 + 0.117 \times 10^{ - 3} (P - 1)$$ for a temperature of 700°K and a plagioclase composition of An25. By combining this equation with equations describing equilibria between graphite and gas species in the system C-H-O, the following partial pressures: \(P{\text{H}}_2 {\text{O}} = 2572{\text{b, }}P{\text{CO}}_2 = 3162{\text{b, }}P{\text{H}}_2 = 2.5{\text{b, }}P{\text{CH}}_4 = 52.5{\text{b, }}P{\text{CO}} = 11.0{\text{b}}\) are obtained for \(f{\text{O}}_2 = 10^{ - 26}\) . If total pressure equals fluid pressure, then the total pressure during metamorphism was approximately 6 kb. The total fluid pressure calculated is extremely sensitive to the value of \(f{\text{O}}_2\) chosen.  相似文献   

2.
The activity-composition relations for calcium-rich and calcium-poor amphiboles are calculated from the composition of coexisting cummingtonite-hornblende pairs from a suite of New Zealand rhyolites. The activities are formulated in terms of site occupancies and the regular solution model is used to represent non-ideal mixing of the cations on each site. The regular solution parameters for each site are calculated from the compositions of the coexisting amphiboles. The resulting activity-composition relations are used to calibrate the standard Gibbs energy change for the reaction $${\text{7MgSiO}}_{\text{3}} {\text{ + SiO}}_{\text{2}} {\text{ + H}}_{\text{2}} {\text{O = Mg}}_{\text{7}} {\text{Si}}_{\text{8}} {\text{O}}_{{\text{22}}} {\text{(OH)}}_{\text{2}} $$ assuming that the lowest temperature rhyolites in this suite crystallised at \(P_{{\text{H}}_2 {\text{O}}} = P_{{\text{total}}} \)   相似文献   

3.
Theoretical and practical considerations are combined to place limits on the iron content of an FePt alloy that is in equilibrium with silicate melt, olivine and a gas phase of known \(f_{{\text{O}}_{\text{2}} }\) . Equilibrium constants are calculated for the reactions: (1) $$2{\text{Fe}}^{\text{o}} + {\text{SiO}}_{\text{2}} + {\text{O}}_{\text{2}} \rightleftharpoons {\text{Fe}}_{\text{2}} {\text{SiO}}_{\text{4}}$$ (2) $${\text{Fe}}^{\text{o}} + \frac{1}{2}{\text{O}}_{\text{2}} \rightleftharpoons {\text{FeO}}$$ . These equilibria may be used to choose an appropriate iron activity for the FePt alloy of an experiment. The temperature dependence of the equilibrium constants is calculated from experimental data. The Gibbs free energy of reaction (1) obtained using thermochemical data is in close agreement with ΔGrxn calculated from the experimental data. Reaction (1) has the advantage that it is independent of the Fe2+/Fe3+ ratio of the melt, but is limited to applications where olivine is a crystallizing phase and requires a formulation for \(a_{{\text{SiO}}_{\text{2}} }^{{\text{liq}}}\) . Reaction (2) uses an empirical approximation for the FeO/Fe2O3 ratio of the liquid, and is independent of olivine saturation. However, it requires a formulation for a FeO liq . Either equilibrium constant may be used to calculate the appropriate FePt alloy in equilibrium with a silicate melt. If experiments are conducted at an \(f_{{\text{O}}_{\text{2}} }\) parallel that of a buffer assemblage, a small range of FePt alloys may be used over a large temperature interval. For example, an alloy containing from 6 % to 9 % Fe by weight is in equilibrium with olivine-saturated tholeiites and komatiites at the quartzfayalite-magnetite buffer over the temperature interval 1,400° C to 1,100° C. Lunar basalt liquids in equilibrium with olivine at 1/2 log unit below the iron-wüstite buffer require an FePt alloy that contains 30–50 wt. % iron over a similar temperature interval.  相似文献   

4.
Stoichiometric mixtures of tremolite and dolomite were heated to 50° C above equilibrium temperatures to form forsterite and calcite. The pressure of the CO2-H2O fluid was 5 Kb and \(X_{{\text{CO}}_{\text{2}} }\) varied from 0.1 to 0.6. The extent of the conversion was determined by the amount of CO2 produced. The resulting mixtures of unreacted tremolite and dolomite and of newly-formed forsterite and calcite were examined with a scanning electron microscope. All tremolite and dolomite grains showed obvious signs of dissolution. At fluid compositions with \(X_{{\text{CO}}_{\text{2}} }\) less than about 0.4, the forsterite and calcite crystals are randomly distributed throughout the charges, indicating that surfaces of the reactants are not a controlling factor with respect to the sites of nucleation of the products. A change is observed when \(X_{{\text{CO}}_{\text{2}} }\) is greater than about 0.4; the forsterite and calcite crystals now nucleate and grow at the surface of the dolomite grains, thus indicating a change in mechanism at medium CO2 concentrations. As the reaction progresses, the dolomite grains become more and more surrounded by forsterite and calcite, finally forming armoured relics of dolomite. Under experimental conditions this characteristic texture can only be formed if the CO2-concentration is greater than about 40 mole %. These findings make it possible to estimate the CO2-concentration from the texture of the dolomite+tremolite+forsterite+calcite assemblage. The results suggest a dissolution-precipitation mechanism for the reaction investigated. In a simplified form it consists of the following 4 steps:
  1. Dissolution of the reactants tremolite and dolomite.
  2. Diffusion of the dissolved constituents in the fluid.
  3. Heterogeneous nucleation of the product minerals.
  4. Growth of forsterite and calcite from the fluid.
Two possible explanations are discussed for the development of the amoured texture at \(X_{{\text{CO}}_{\text{2}} }\) above 0.4. The first is based upon the assumption that dolomite has a lower rate of dissolution than tremolite at high \(X_{{\text{CO}}_{\text{2}} }\) values resulting in preferential calcite and forsterite nucleation and growth on the dolomite surface. An alternative explanation is the formation of a raised CO2 concentration around the dolomite grains at high \(X_{{\text{CO}}_{\text{2}} }\) values, leading to product precipitation on the dolomite crystals.  相似文献   

5.
A new determination of the equilibrium reaction: $$\begin{gathered} 2{\text{ Mg}}_{\text{2}} [{\text{SiO}}_{\text{4}} ] + 3{\text{ H}}_{\text{2}} {\text{O}} \rightleftharpoons {\text{1 Mg}}_{\text{3}} [({\text{OH)}}_{\text{4}} |{\text{Si}}_{\text{2}} {\text{O}}_{\text{5}} ] + 1{\text{ Mg(OH)}}_{\text{2}} \hfill \\ \hfill \\ {\text{ forsterite serpentine brucite}} \hfill \\ \end{gathered} $$ yielded equilibrium temperatures which lie (at identical H2O-pressures) about 60° C lower than all previously published data (Bowen and Tuttle, 1949; Yoder, 1952; Kitahara et al., 1966; Kitahara and Kennedy, 1967). It has been shown that the above authors have determined not the stable equilibrium curve but instead a metastable “synthesis boundary”. The actual (stable) equilibrium curve is located at 0,5 kb and 350° C 2,0 kb and 380° C 3,5 kb and 400° C 5,0 kb and 420° C 6,5 kb and 430° C.  相似文献   

6.
In a regional metamorphic terrain where six isograds have been mapped based on mineral reactions that are observed in metacarbonate rocks, the P-T conditions and fugacities of CO2 and H2O during metamorphism were quantified by calculations involving actual mineral compositions and experimental data. Pressure during metamorphism was near 3,500 bars. Metamorphic temperatures ranged from 380° C (biotite-chlorite isograd) to 520° C (diopside isograd). \(f_{{\text{CO}}_{\text{2}} }\) and \(f_{{\text{CO}}_{\text{2}} }\) / \(f_{{\text{H}}_{\text{2}} {\text{O}}}\) in general is higher in metacarbonate rocks below the zoisite isograd than in those above the zoisite isograd. Calculated \(f_{{\text{CO}}_{\text{2}} }\) and \(f_{{\text{H}}_{\text{2}} {\text{O}}}\) are consistent with carbonate rocks above the zoisite isograd having equilibrated during metamorphism with a bulk supercritical fluid in which \(P_{{\text{CO}}_{\text{2}} }\) + \(P_{{\text{H}}_{\text{2}} {\text{O}}}\) = P total. Calculations indicate that below the zoisite isograd, however, \(P_{{\text{CO}}_{\text{2}} }\) + \(P_{{\text{H}}_{\text{2}} {\text{O}}}\) was less than Ptotal, and that this condition is not due to the presence of significant amounts of species other than CO2 and H2O in the system C-O-H-S. Calculated \(P_{{\text{CO}}_{\text{2}} }\) /( \(P_{{\text{CO}}_{\text{2}} }\) + \(P_{{\text{H}}_{\text{2}} {\text{O}}}\) ) is low (0.06–0.32) above the zoisite isograd. The differences in conditions above and below the zoisite isograd may indicate that the formation of zoisite records the introduction of a bulk supercritical H2O-rich fluid into the metacarbonates. The results of the study indicate that \(f_{{\text{CO}}_{\text{2}} }\) and \(f_{{\text{H}}_{\text{2}} {\text{O}}}\) are constant on a thin section scale, but that gradients in \(f_{{\text{CO}}_{\text{2}} }\) and \(f_{{\text{H}}_{\text{2}} {\text{O}}}\) existed during metamorphism on both outcrop and regional scales.  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.
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   

9.
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 (1) $$\begin{gathered} {\text{Na(LiAl}}_{\text{2}} {\text{) [Al}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{{\text{10}}} {\text{] (OH)}}_{\text{2}} \hfill \\ {\text{ephesite}} \hfill \\ {\text{ = Na[AlSiO}}_{\text{4}} {\text{] + LiAl[SiO}}_{\text{4}} {\text{] + }}\alpha {\text{ - Al}}_{\text{2}} {\text{O}}_{\text{3}} {\text{ + H}}_{\text{2}} {\text{O}} \hfill \\ {\text{nepheline }}\alpha {\text{ - eucryptite corundum}} \hfill \\ \end{gathered}$$ Five reversal brackets for (1) have been established experimentally in the temperature range 590–750° C, at \(P_{H_2 O}\) between 400 and 2500 bars. The equilibrium constant, K, for this reaction may be expressed as (2) $$log K{\text{ = }}log f_{{\text{H}}_{\text{2}} O}^* = 7.5217 - 4388/T + 0.0234 (P - 1)T$$ where \(f_{H_2 O}^* = f_{H_2 O} (P,T)/f_{H_2 O}^0\) (1,T), with T given in degrees K, and P in bars. Combining these experimental data with known thermodynamic properties of the decomposition products in (1), the following standard state (1 bar, 298.15 K) thermodynamic data for ephesite were calculated: H f,298.15 0 =-6237372 J/mol, S 298.15 0 =300.455 J/K·mol, G 298.15 0 =-5851994 J/mol, and V 298.15 0 =13.1468 J/bar·mol.  相似文献   

10.
An empirically derived Redlich-Kwong type of equation of state (ERK) is proposed for H2O, expressing a, the term related to the attraction between the molecules, as a pressure-independent function of temperature, and b, the covolume, as a temperature-independent function of pressure. The coefficients of a(T) and b(P) were derived by least squares non-linear regression, using P-V-T data given by Burnham et al. (1969b) and Rice and Walsh (1957) in conjunction with more precise recent data obtained by Tanishita et al. (1976), Hilbert (1979) and Schmidt (1979): $$a(T) = 1.616 x 10^8 - 4.989 x 10^4 T - 7.358 x 10^9 T^{ - 1} $$ and $$ = \frac{{1 + 3.4505x 10^{--- 4} P + 3.8980x 10^{--- 9} P^2 - 2.7756x 10^{--- 15} P^3 }}{{6.3944x 10^{--- 2} + 2.3776x 10^{--- 5} + 4.5717x 10^{--- 10} P^2 }}$$ , where T is expressed in Kelvin and P in bars. The ERK works very well at upper mantle conditions, at least up to 200 kbar and 1,000 °C. At subcritical conditions and those somewhat above the critical point, it still reproduces the molar Gibbs energy, \(\tilde G_{{\text{H}}_{\text{2}} {\text{O}}} \) , with a maximum deviation of 400 joules. Thus, for the purpose of calculation of geologically interesting heterogeneous equilibria, it predicts the thermodynamic properties of H2O well enough. The values of molar volume, \(\tilde V_{{\text{H}}_{\text{2}} {\text{O}}} \) , and \(\tilde G_{{\text{H}}_{\text{2}} {\text{O}}} \) are tabulated in the appendix over a considerable P-T range. A FORTRAN program generating these functions as well as a FORTRAN subroutine for calculating the fugacity values, \(f_{{\text{H}}_{\text{2}} {\text{O}}} \) for incorporation into existing programs, are available upon request.  相似文献   

11.
The equilibrium curve for the reaction 3 dolomite + 1 K-feldspar + 1 H2O=1 phlogopite + 3 calcite + 3 CO2 was determined experimentally at a total gas pressure of 2000 bars using two different methods.
  1. In the first case water alone was added to the reactants. The CO2 component of the gas phase was producted solely by the reaction under favourable P-T conditions. This manner of carrying out the reaction is called the “water method”. With this method sufficient time must be allowed for the gas phase to attain a constant composition (see Fig. 1). Reverse reactions were carried out using reaction products of the forward reaction.
  2. In the second case silver oxalate + water were added to the reactants. Breakdown of the silver oxalate leads to formation of a CO2-H2O gasphase of definite composition. At constant temperature and gas pressure the \(X_{{\text{CO}}_{\text{2}} } \) determines whether the reaction products will be phlogopite + calcite or dolomite + K-feldspar. In this case it is not necessary to wait for equilibrium to be attained. This method is abbreviated the “oxalate method”. Reactants for reverse reactions are not identical with the products of the forward reaction.
At high temperatures the results of the two different methods agree well (see Tables 1 and 2). Equilibrium was attained in one case at 490° C and \(X_{{\text{CO}}_{\text{2}} } \) of approximately 0.77, and in the other case at 520° C and \(X_{{\text{CO}}_{\text{2}} } \) of 0.90. At lower temperatures there are considerable differences in the results. With the water method an \(X_{{\text{CO}}_{\text{2}} } \) of about 0.25 was reached at 450° C. With the oxalate method dolomite K-feldspar and water still react with each other at even higher \(X_{{\text{CO}}_{\text{2}} } \) values. Phlogopite, calcite and CO2 are formed together with metastable talc. There are no criteria to indicate which of the methods is the correct one at lower temperatures and in Fig. 2, therefore, both equilibrium curves are plotted.  相似文献   

12.
A great wealth of analytical data for fluid inclusions in minerals indicate that the major species of gases in fluid inclusions are H2O, CO2, CO, CH4, H2 and O2. Three basic chemical reactions are supposed to prevail in rock-forming and ore-forming fluids: $$\begin{gathered} H_2 + 1/2{\text{ O}}_{\text{2}} = H_2 O, \hfill \\ CO + 1/2{\text{ O}}_{\text{2}} = CO_2 , \hfill \\ CH_4 + 2{\text{O}}_{\text{2}} = CO_2 + 2H_2 O, \hfill \\ \end{gathered} $$ and equilibria are reached among them. \(\lg f_{O_2 } - T,{\text{ }}\lg f_{CO_2 } - T\) and Eh-T charts for petrogenesis and minerogenesis in the supercritical state have been plotted under different pressures. On the basis of these charts \(f_{O^2 } ,{\text{ }}f_{CO_2 } \) , Eh, equilibrium temperature and equilibrium pressure can be readily calculated. In this paper some examples are presented to show their successful application in the study of the ore-forming environments of ore deposits.  相似文献   

13.
P, T, \(X_{{\text{CO}}_{\text{2}} }\) relations of gehlenite, anorthite, grossularite, wollastonite, corundum and calcite have been determined experimentally at P f =1 and 4 kb. Using synthetic starting minerals the following reactions have been demonstrated reversibly
  1. 2 anorthite+3 calcite=gehlenite+grossularite+3 CO2.
  2. anorthite+corundum+3 calcite=2 gehlenite+3 CO2.
  3. 3anorthite+3 calcite=2 grossularite+corundum+3CO2.
  4. grossularite+2 corundum+3 calcite=3 gehlenite+3 CO2.
  5. anorthite+2 calcite=gehlenite+wollastonite+2CO2.
  6. anorthite+wollastonite+calcite=grossularite+CO2.
  7. grossularite+calcite=gehlenite+2 wollastonite+CO2.
In the T, \(X_{{\text{CO}}_{\text{2}} }\) diagram at P f =1 kb two isobaric invariant points have been located at 770±10°C, \(X_{{\text{CO}}_{\text{2}} }\) =0.27 and at 840±10°C, \(X_{{\text{CO}}_{\text{2}} }\) =0.55. Formation of gehlenite from low temperature assemblages according to (4) and (2) takes place at 1 kb and 715–855° C, \(X_{{\text{CO}}_{\text{2}} }\) =0.1–1.0. In agreement with experimental results the formation of gehlenite in natural metamorphic rocks is restricted to shallow, high temperature contact aureoles.  相似文献   

14.
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   

15.
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   

16.
This study presents accurate and precise iron isotopic data for 16 co-magmatic rocks and 6 pyroxene–magnetite pairs from the classic, tholeiitic Red Hill sill in southern Tasmania. The intrusion exhibits a vertical continuum of compositions created by in situ fractional crystallisation of a single injection of magma in a closed igneous system and, as such, constitutes a natural laboratory amenable to determining the causes of Fe isotope fractionation in magmatic rocks. Early fractionation of pyroxenes and plagioclase, under conditions closed to oxygen exchange, gives rise to an iron enrichment trend and an increase in $ f_{{{\text{O}}_{2} }} $ of the melt relative to the Fayalite–Magnetite–Quartz (FMQ) buffer. Enrichment in Fe3+/ΣFemelt is mirrored by δ57Fe, where VIFe2+-bearing pyroxenes partition 57Fe-depleted iron, defining an equilibrium pyroxene-melt fractionation factor of $ \Updelta^{57} {\text{Fe}}_{{{\text{px}} - {\text{melt}}}} \le - 0.25\,\permille \times 10^{6} /T^{2} $ . Upon magnetite saturation, the $ f_{{{\text{O}}_{2} }} $ and δ57Fe of the melt fall, commensurate with the sequestration of the oxidised, 57Fe-enriched iron into magnetite, quantified as $ \Updelta^{57} {\text{Fe}}_{{{\text{mtn}} - {\text{melt}}}} = + 0.20\,\permille \times 10^{6} /T^{2} $ . Pyroxene–magnetite pairs reveal an equilibrium fractionation factor of $ \Updelta^{57} {\text{Fe}}_{{{\text{mtn}} - {\text{px}}}} \approx + 0.30\,\permille $ at 900–1,000?°C. Iron isotopes in differentiated magmas suggest that they may act as an indicator of their oxidation state and tectonic setting.  相似文献   

17.
Oxygen Fugacity measurements were carried out on chromites from the Eastern Bushveld Complex (Maandagshoek) and are compared with former measurements on chromites from the western Bushveld Complex (Zwartkop Chrome Mine). These results together with those of Hill and Roeder (1974) yield the following conditions of formation for the massive chromitite layers: Western Bushveld Complex (Zwartkop Chrome Mine) $$\begin{gathered} Layer{\text{ }}T(^\circ C) p_{O_2 } (atm) \hfill \\ LG3{\text{ 1160}} - {\text{1234 10}}^{ - {\text{5}}} - 10^{ - 7.6} \hfill \\ LG4{\text{ 1175}} - {\text{1200 10}}^{ - 6.35} - 10^{ - 7.20} \hfill \\ LG6{\text{ 1162}} - {\text{1207 10}}^{ - 6.20} - 10^{ - 7.50} \hfill \\ \hfill \\ \end{gathered} $$ Eastern Bushveld Complex (Farm Maandagshoek) $$\begin{gathered} {\text{LXI 1115}} - {\text{1150 10}}^{ - 7.80} - 10^{ - 8.80} \hfill \\ ( = {\text{Steelpoort Seam)}} \hfill \\ {\text{LX 1125 10}}^{ - 8.25} \hfill \\ {\text{V 1120 10}}^{ - 8.55} \hfill \\ {\text{LII 1120 10}}^{ - 8.0} - 10^{ - 8.60} \hfill \\ \end{gathered} $$ The comparison of the data shows, that the chronitite layers within each particular sequence were formed under approximately identicalp o 2- andT-conditions. The chromites from the western Bushveld Complex, however, were formed at higher temperatures and higher oxygen fugacities than the chromites from the eastern Bushveld Complex. Fromp o 2-T-curves of disseminated chromites and the temperatures derived above, the following conditions of formation for the host rocks were obtained: Western Bushveld Complex $$T = 1200^\circ {\text{C; }}p_{{\text{o}}_{\text{2}} } = 10^{ - 7.25} - 10^{ - 7.50} $$ Eastern Bushveld Complex $$T = 1125^\circ {\text{C; }}p_{{\text{o}}_{\text{2}} } = 10^{ - 8.50} - 10^{ - 9.25} $$ Consequently, the host rocks in the Zwartkop-Chrome-Mine, were formed under higher temperatures and higher oxygen fugacities than the host rocks at Maandagshoek. The rock sequence in the Zwartkop-Chrome-Mine therefore originated in an earlier stage of the differentiation of the Bushveld magma. Comparison of the chromites from the host rocks with the chromites from massive layers supports Ulmer's (1969) thesis that an increase of the oxygen fugacity is responsible for the formation of massive chromitite layers. The values in this investigation show that increases of only about 0.5–1.0 log units are necessary to enhance chromitite layer formation.  相似文献   

18.
Natural springs have been reliable sources of domestic water and have allowed for the development of recreational facilities and resorts in the Central Appalachians. The structural history of this area is complex and it is unknown whether these natural springs receive significant recharge from modern precipitation or whether they discharge old water recharged over geological times scales. The main objective of this study was to use stable isotopes of water ( $\delta^{18} {\text{O}}_{{{\text{H}}_{2} {\text{O}}}}$ and $\delta^{2} {\text{H}}_{{{\text{H}}_{2} {\text{O}}}}$ ), dissolved inorganic carbon ( $\delta^{13} {\text{C}}_{\text{DIC}}$ ) and dissolved sulfate ( $\delta^{34} {\text{S}}_{{{\text{SO}}_{4} }}$ and $\delta^{18} {\text{O}}_{{{\text{SO}}_{4} }}$ ) to delineate sources of water, carbon and sulfur in several natural springs of the region. Our preliminary isotope data indicate that all springs are being recharged by modern precipitation. The oxygen isotope composition indicates that waters in thermal springs did not encounter the high temperatures required for O isotope exchange between the water and silicate/carbonate minerals, and/or the residence time of water in the aquifers was short due to high flow rates. The carbon isotopic composition of dissolved inorganic carbon and sulfur/oxygen isotopic composition of dissolved sulfate provide evidence of low-temperature water–rock interactions and various biogeochemical transformations these waters have undergone along their flow path.  相似文献   

19.
The ferric-ferrous ratio of natural silicate liquids equilibrated in air   总被引：1，自引：1，他引：1  
Results of chemical analyses of glasses produced in 46 melting experiments in air at 1,350° C and 1,450° C on rocks ranging in composition from nephelinite to rhyolite have been combined with other published data to obtain an empirical equation relating in \((X_{{\text{Fe}}_{\text{2}} {\text{O}}_{\text{3}} }^{{\text{liq}}} /X_{{\text{FeO}}}^{{\text{liq}}} )\) to T, \(\ln f_{{\text{O}}_{\text{2}} } \) and bulk composition. The whole set of experimental data range over 1,200–1,450° C and oxygen fugacities of 10?9.00 to 10?0.69 bars, respectively. The standard errors of temperature and \(\log _{10} f_{{\text{O}}_{\text{2}} } \) predictions from this equation are 52° C and 0.5 units, respectively, for 186 experiments.  相似文献   

20.
The stability relations between cordierite and almandite in rocks, having a composition of CaO poor argillaceous rocks, were experimentally investigated. The starting material consisted of a mixture of chlorite, muscovite, and quartz. Systems with widely varying Fe2+/Fe2++Mg ratios were investigated by using two different chlorites, thuringite or ripidolite, in the starting mixture. Cordierite is formed according to the following reaction: $${\text{Chlorite + muscovite + quartz}} \rightleftharpoons {\text{cordierite + biotite + Al}}_{\text{2}} {\text{SiO}}_{\text{5}} + {\text{H}}_{\text{2}} {\text{O}}$$ . At low pressures this reaction characterizes the facies boundary between the albite-epidotehornfels facies and the hornblende-hornfels facies, at medium pressures the beginning of the cordierite-amphibolite facies. Experiments were carried out reversibly and gave the following equilibrium data: 505±10°C at 500 bars H2O pressure, 513±10°C at 1000 bars H2O pressure, 527±10°C at 2000 bars H2O pressure, and 557±10°C at 4000 bars H2O pressure. These equilibrium data are valid for the Fe-rich starting material, using thuringite as the chlorite, as well as for the Mg-rich starting mixture with ripidolite. At 6000 bars the equilibrium temperature for the Mg-rich mixture is 587±10°C. In the Fe-rich mixture almandite was formed instead of cordierite at 6000 bars. The following reaction was observed: $${\text{Thuringite + muscovite + quartz}} \rightleftharpoons {\text{almandite + biotite + Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{ + H}}_{\text{2}} {\text{O}}$$ . Experiments with the Fe-rich mixture, containing Fe2+/Fe2++Mg in the ratio 8∶10, yielded three stability fields in a P,T-diagram (Fig.1):
  1. Above 600°C/5.25 kb and 700°C/6.5 kb almandite+biotite+Al2SiO5 coexist stably, cordierite being unstable.
  2. The field, in which almandite, biotite and Al2SiO5 are stable together with cordierite, is restricted by two curves, passing through the following points:
    1. 625°C/5.5 kb and 700°C/6.5 kb,
    2. 625°C/5.5 kb and 700°C/4.0 kb.
  3. At conditions below curves 1 and 2b, cordierite, biotite, and Al2SiO5 are formed, but no garnet.
An appreciable MnO-content in the system lowers the pressures needed for the formation of almandite garnet, but the quantitative influence of the spessartite-component on the formation of almandite could not yet be determined. the Mg-rich system with Fe2+/Fe2++Mg=0.4 garnet did not form at pressures up to 7 kb in the temperature range investigated. Experiments at unspecified higher pressures (in a simple squeezer-type apparatus) yielded the reaction: $${\text{Ripidolite + muscovite + quartz}} \rightleftharpoons {\text{almandite + biotite + Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{ + H}}_{\text{2}} {\text{O}}$$ . Further experiments are needed to determine the equilibrium data. The occurence of garnet in metamorphic rocks is discussed in the light of the experimental results.  相似文献   

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