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1.
Fractionations of carbon and oxygen isotopes and magnesium between coexisting dolomite and calcite have been determined for marbles and calcareous schists of a wide variety of metamorphic environments from Vermont and the Grenville Province of Ontario. Concordant equilibrium fractionations are given by 83% of the samples. Calibration of the isotopic thermometers using the Mg-calcite solvus thermometer gave in the temperature range: 650°>T°>100°C $$ \begin{gathered} 1,000\ln \alpha _{D - Ct}^{O^{18} } = 0.45 (10^6 T^{ - 2} ) - 0.40 \hfill \\ 1,000\ln \alpha _{D - Ct}^{O^{18} } = 0.18 (10^6 T^{ - 2} ) + 0.17. \hfill \\ \end{gathered} $$ These isotopic fractionation expressions differ significantly from the experimentally derived relations, including the dolomite-Mg-calcite C13 partial exchange experiments of this study. Temperature ranges obtained for the metamorphic zones of Vermont are: chlorite zone, 210° to 295° C; biotite zone, 255° to 400° C; staurolite-kyanite zone, 110° to 550° C. In amphibolite-facies rocks the quenched partition relations can be complex. The temperature of quench or recrystallization may be as large as 400° C below the inferred metamorphic maximum. Oxygen isotope disequilibrium in high grade rocks, particularly from the Chester dome area, Vermont, is characterized by large negative δO D 18 –δO Ct 18 values. The size of the equilibrium exchange system for carbon and oxygen isotopes and magnesium is small, less than a few inches across the inferred relict bedding. This is attributed to the lack of a mobile pore fluid except in systems undergoing decarbonation. C13/C12 ratios in Grenville and Vermont marbles and O18/O16 ratios in Grenville and greenschist-facies Vermont carbonates span the range of ancient limestones. Staurolite-kyanite zone calcareous schists and marbles from the Chester dome area, Vermont are depleted in O18(δO18=12 to 20‰) due to equilibrium or disequilibrium decarbonation and some partial exchange. Extrapolation of the dolomite-calcite fractionation expressions to 20° C indicates that dolomite is enriched in O18 by about 4.9‰ and in C13 by about 2.4‰.  相似文献   

2.
In the present work we studied Mg-ilmenite megacrysts from the Arkhangelsk kimberlites (the Kepino kimberlite field and mantle xenoliths from the Grib pipe). On the basis of isotopic (Rb/Sr, Sm/Nd, δ18O) and trace-element data we argue that studied Mg-ilmenite megacrysts have a genetic relation to the “protokimberlitic” magma, which was parental to the host kimberlites. Rb-Sr ages measured on phlogopite from ilmenite-clinopyroxenite xenoliths and the host Grib kimberlite overlap within the error (384 Ma and 372 ± 8 Ma, respectively; Shevchenko et al., 2004) with our estimation of the Kotuga kimberlite emplacement (378 ± 25 Ma). Sr and Nd isotopic compositions of megacrysts are close to the isotopic composition of host kimberlites (Mg-ilmenites from kimberlites have 87Sr/86Sr(t = 384) = 0.7050–0.7063, ?Nd(t = 384) = + 1.7, +1.8, ilmenite from ilmenite-garnet clinopyroxenite xenolith has 87Sr/86St(t = 384) = 0.7049, ?Nd(t = 384) = +3.5). Oxygen isotopic composition of ilmenites (δ18O = +3.8–+4.5‰) is relatively “light” in comparison with the values for mantle minerals (δ18O = +5–+6‰). Taking into account ilmenite-melt isotope fractionation, these values of δ18O indicate that ilmenites could crystallize from the “protokimberlitic” melt. Temperatures and redox conditions during the formation of ilmenite reaction rims were estimated using ilmenite-rutile and titanomagnetite-ilmenite thermo-oxybarometers. New minerals within the rims crystallized at increasing oxygen fugacity and decreasing temperature. Spinels precipitated during the interaction of ilmenite with kimberlitic melt at T = 1000–1100°C and oxygen fugacity $\Delta \log f_{O_2 }$ [QFM] ≈ 1. Rims comprised with rutile and titanomagnetite crystallized at T ≈ 1100°C, $\Delta \log f_{O_2 }$ [NNO] ≈ 4 and T = 600–613°C, $\Delta \log f_{O_2 }$ [QFM] ≈ 3.7, respectively. Rutile lamellae within ilmenite grains from clinopyroxenitic xenolith were formed T ≥ 1000–1100°C and oxygen fugacity $\Delta \log f_{O_2 }$ [NNO] = ?3.7. Since the pressure of clinopyroxene formation from this xenolith was estimated to be 45–53 kbar, redox conditions at 135–212 km depths could be close to $\Delta \log f_{O_2 }$ [NNO] = ?3.7.  相似文献   

3.
Oxygen diffusion in albite has been determined by the integrating (bulk 18O) method between 750° and 450° C, for a P H2O of 2 kb. The original material has a low dislocation density (<106 cm?2), and its lattice diffusion coefficient (D 1), given below, agrees well with previous determinations. A sample was deformed at high temperature and pressure to produce a uniform dislocation density of 5 × 109 cm?2. The diffusion coefficient (D a) for this deformed material, given below, is about 0.5 and 0.7 orders of magnitude larger than D 1 at 700° and 450° C, respectively. This enhancement is believed due to faster diffusion along the cores of dislocations. Assuming a dislocation core radius of 4 Å, the calculated pipe diffusion coefficient (D p), given below, is about 5 orders of magnitude larger than D 1. These results suggest that volume diffusion at metamorphic conditions may be only slightly enhanced by the presence of dislocations. $$\begin{gathered} D_1 = 9.8 \pm 6.9 \times 10^{ - 6} (cm^2 /\sec ) \hfill \\ {\text{ }} \cdot \exp [ - 33.4 \pm 0.6(kcal/mole)/RT] \hfill \\ \end{gathered} $$ $$\begin{gathered} D_a = 7.6 \pm 4.0 \times 10^{ - 6} (cm^2 /\sec ) \hfill \\ {\text{ }} \cdot \exp [ - 30.9 \pm 1.1(kcal/mole)/RT] \hfill \\ \end{gathered} $$ $$\begin{gathered} D_p \approx 1.2 \times 10^{ - 1} (cm^2 /\sec ) \hfill \\ {\text{ }} \cdot \exp [ - 29.8(kcal/mole)/RT]. \hfill \\ \end{gathered} $$   相似文献   

4.
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   

5.
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   

6.
Piemontite- and thulite-bearing assemblages from highly oxidized metapelitic and metacalcareous schists associated with braunite quartzites at Vitali, Andros island, Greece, were chemically investigated. The Mn-rich metasediments are intercalated in a series of metapelitic quartzose schists, marbles, and basic metavolcanites which were affected by a regional metamorphism of the highP/T type (T=400–500° C,P>9 kb) and a later Barrovian-type greenschist metamorphism (T=400–500° C,P～-5–6 kb). Texturally and chemically two generations of piemontite (I and II) can be distinguished which may show complex compositional zoning. Piemontite I coexisted at highP/T conditions with braunite, manganian phengite (alurgite), Mn3+-Mn2+-bearing Na-pyroxene (violan), carbonate, quartz, hollandite, and hematite. Zoned grains generally exhibit a decreasing Mn3+ and an increasing Fe3+ and Al content towards the rim. Chemical compositions of piemontite I range from 2.0 to 32.1 mole % Mn3+, 0 to 25.6 mole % Fe3+, and 60.2 to 81.2 mole % Al. Up to 12.5 mole % Ca on the A(2) site can be substituted by Sr. Piemontites formed in contact or close to braunite (±hematite) attained maximum (Mn3++Fe3+)Al?1 substitution corrresponding to about 33 mole % Mn3++Fe3+ in lowiron compositions and up to about 39 mole % Mn3++ Fe3+ at intermediate Fe3+/(Fe3++Mn3+) ratios. Piemontite II which discontinuously overgrows piemontite I or occurs as separate grains may have been formed by greenschist facies decomposition of manganian Na-pyroxenes according to the reaction: (1) $$\begin{gathered} {\text{Mn}}^{{\text{3 + }}} - Mn^{2 + } - bearing omphacite/chloromelanite \hfill \\ + CO_2 + H_2 O + HCl \pm hermatite \hfill \\ = piemontite + tremolite + albite + chlorite \hfill \\ + calcite + quartz + NaCl \pm O_2 . \hfill \\ \end{gathered} $$ Thulites crystallized in coexistence with Al-rich piemontite II. All thulites analysed are low-Fe3+ manganian orthozoisites with Mntot～-Mn3+ substituting for Al on the M(3) site. Their compositions range from 2.9 to 7.2 mole % Mn3+, 0 to 1.2 mole % Fe3+, and 91.8 to 96.7 mole % Al. Piemontites II in thulite-bearing assemblages range from 5.8 to 15.9 mole % Mn3+, 0 to 3.7 mole % Fe3+, and 83.7 to 93.6 mole % Al. By contrast, piemontites II in thulite-free assemblages are similarly enriched in Mn3+ + Fe3+ — and partially in Sr2+ — as core compositions of piemontite I (21.1 to 29.6 mole % Mn3+, 2.0 to 16.5 mole % Fe3+, 60.6 to 68.4 mole % Al, 0 to 29.4 mole % Sr in the A(2) site). The analytical data presented in this paper document for the first time a continuous low-Fe3+ piemontite solid solution series from 5.8 to 32.1 mole % Mn3+. Aluminous piemontite II is enriched by about 3 mole % Mn3++Fe3+ relative to coexisting thulite in Fe3+-poor samples and by about 6 mole % Mn3++Fe3+ in more Fe3+-rich samples. Mineral pairs from different samples form a continuous compositional loop. Compositional shift of mineral pairs is attributed to the effect of a variable fluid composition at constantP fluid andT on the continuous reaction: (2) $$\begin{gathered} piemontite + CO_2 \hfill \\ = thulite + calcite + quartz \hfill \\ + Mn^{2 + } Ca_{ - 1} [calcite] + H{_2} O + O{_2} \hfill \\ \end{gathered} $$ Further evidence for a variable \(x_{H_2 O} \) and/or \(f_{O_2 } \) possibly resulting from fluid infiltration and local buffering during the greenschist metamorphism is derived from the local decomposition of piemontite, braunite, and rutile to form spessartine, calcite, titanite, and hematite by the reactions: (3) $$\begin{gathered} piemontite + braunite + CO_2 \hfill \\ = sperssartine + calcite + quartz \pm hermatite \hfill \\ + H{_2} O + O{_2} \hfill \\ \end{gathered} $$ and more rarely: (4) $$\begin{gathered} piemontite + quartz + rutile + braunite \hfill \\ = spessartine + titanite + hematite + H{_2} O + O{_2} . \hfill \\ \end{gathered} $$   相似文献   

7.
Geochemical potential field is defined as the scope within the earth’s space where a given component in a certain phase of a certain material system is acted upon by a diffusion force, depending on its spatial coordinatesX, Y andZ. The three coordinates follow the relations: $$NF_{ix} = - \frac{{\partial \mu }}{{\partial x}}, NF_{iy} = - \frac{{\partial \mu }}{{\partial y}}, NF_{iz} = - \frac{{\partial \mu }}{{\partial z}}$$ The characteristics of such a field can be summarized as: (1) The summation of geochemical potentials related to the coordinatesX, Y, Z, or pseudo-velocity head, pseudo-pressure head and pseudo-potential head of a certain component in the earth is a constant as given by $$\mu _x + \mu _y + \mu _z = c$$ or $$\mu _{x2} + \mu _{y2} + \mu _{z2} = \mu _{x1} + \mu _{y1} + \mu _{z1} $$ Derived from these relations is the principle of geochemical potential conservation. The following relations have the same physical significance: $$\mu _k + \mu _u + \mu _p = c$$ or $$\mu _{k2} + \mu _{u2} + \mu _{p2} = \mu _{k1} + \mu _{u1} + \mu _{p1} $$ (2) Geochemical potential field is a vector field quantified by geochemical field intensity which is defined as the diffusion force applied to one molecular volume (or one atomic volume) of a certain component moving from its higher concentration phase to lower concentration phase. The geochemical potential field intensity is given by $$\begin{gathered} E = - grad\mu \hfill \\ E = \frac{{RT}}{x}i + \frac{{RT}}{y}j + \frac{{RT}}{z}K \hfill \\ \end{gathered} $$ The present theory has been inferred to interpret the mechanism of formation of some tungsten ore deposits in China.  相似文献   

8.
Oilfield brines (produced water) are produced as a waste product daily at the gathering centers (GCs) in Kuwait oilfields. The geochemical evolution of the water produced at the GC (fresh brine) to stagnant pit water (evaporate) has been investigated in the northern fields of Kuwait, and a model is presented showing time-dependent variations. Kuwait oilfield brines are globally similar to others in other large sedimentary basins (USA, Canada), but modifications have occurred due to seawater injection practices performed episodically during the oil extraction process. Brine water chemistry changes from generally average brine chemistry (based on cations and anions) to saturated mixture of seawater, oilfield brine, and anthropogenic chemical pollutants. The objective of this study was to harmonize the database of brine waters in terms of regional identity by comparison with oilfield brines elsewhere, identify water–rock interaction, and statistically treat daily recordings from the pits in order to identify injection peaks and troughs. Laboratory analysis of major and minor cations and anions from the Rawdatayn samples gave the following concentration ranges in parts per million (ppm): (Na+, 11,698–203,977), (Ca2+, 2,216–98,514), (Mg2+, 1,602–28,885), (K+, 1,528–16,573), (Sr2+, 70–502), (Ba2+, 0.01–18.04), (Fe2+, 0.01–8.93), (Li+, 0.09–6.48), (Si2+, 0.00–13.18), (B3+, 0.05–37.45), (SO 4 2+ , 330–3100). For the Sabriyah oilfield samples, the major and minor cations and anions concentration ranges in ppm are: (Na+, 9,807–274,947), (Ca2+, 2,555–77,992), (Mg2+, 1,415–28,183), (K+, 764–19,201), (Sr2+, 77.84–641), (Ba2+, 0.15–6.76), (Fe2+, 0.016–38.88), (Li+, 0.05–6.83), (Si2+, 0.0195–16.84), (B3+, 7.17–55.33), (SO 4 2+ , 44,812–135,264). The stable isotopic analysis of five samples indicates normal trends in oxygen and hydrogen isotopes that classify the waters as “connate” which follow an evaporation trend. Carbon isotopic signatures are normal for hydrocarbon fields and average out around GC15, δ18O‰?=?1.4, δD‰?=??10, δ13C‰?=??3.6; while for GC23, δ18O‰?=?2.3, δD‰?=??4, δ13C‰?=??2.5; for GC25, δ18O‰?=??2.0, δD‰?=??14, δ13C‰?=??4.6; for pit1, δ18O‰?=?2.3, δD‰?=??5, δ13C‰?=??18.3; and for pit 2, δ18O‰?=?2.5, δD‰?=??4, δ13C‰?=??17.8. Carbon isotope average values for all brine samples from the GCs is?=??56 which falls within normal hydrocarbon formation water category. Data spikes coincide with injection periods at the following times (A: May–Jun, 2006), (B: Sep–Oct, 2006), (C: Jan–Feb, 2007), (D: Mar, 2007), (E: May–Jun, 2007), (F: Feb, 2006), (G: Mar–Apr, 2006) and, subsequently the decay to “normal” brine occurs over a period of several weeks. The database was large enough to apply a principal component statistical analysis (PCA). PCA and geo-statistical techniques reveal several distinct population groups. The main chemical groups in the data are as follows: plateau, spike groups, and pit evaporation group. The spike periods correlate closely with seawater injection periods (Jan–Feb, Mar–Apr, May–Jun, and Sep–Oct). The pit chemistry reveals exceptionally high evaporation processes coinciding with summer peak temperature. PCA results show distinct groupings centered around the major elements reminiscent of other oilfields, but with the added evaporation trend strongly enhanced.  相似文献   

9.
Oxygen isotope fractionation between rutile and water   总被引：1，自引：0，他引：1  
Synthetic rutile-water fractionations (1000 ln α) at 775, 675, and 575° C were found to be ?2.8, ?3.5, and ?4.8, respectively. Partial exchange experiments with natural rutile at 575° C and with synthetic rutile at 475° C failed to yield reliable fractionations. Isotopic fractionation within the range 575–775° C may be expressed as follows: 1 $$1000\ln \alpha ({\rm T}i{\rm O}_{2 } - H_2 O) = - 4.1 \frac{{10^6 }}{{T_{k^2 } }} + 0.96$$ . Combined with previously determined quartz-water fractionations, the above data permit calibration of the quartz-rutile geothermometer: 1 $$1000\ln \alpha ({\text{S}}i{\rm O}_{2 } - Ti{\rm O}_{2 } ) = 6.6 \frac{{10^6 }}{{T_{k^2 } }} - 2.9$$ . When applied to B-type eclogites from Europe, as an example, the latter equation yields a mean equilibration temperature of 565° C.  相似文献   

10.
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   

11.
The complexation between gold and silica was experimentally, confirmed and calibrated at 200 °C: $$\begin{gathered} Au^ + + H_3 SiO_4^ - \rightleftharpoons AuH_3 SiO_4^0 \hfill \\ \log K_{(200^\circ C)} = 19.26 \pm 0.4 \hfill \\ \end{gathered} $$ Thermodynamic calculations show that AuH3SiO 4 0 would be far more abundant than AuCl 2 ? under physicochemical conditions of geological interest, suggesting that silica is much more important than chloride as ligands for gold transport. In systems containing both sulfur and silica, AuH3SiO 4 0 would be increasingly more important than Au (HS) 2 ? as the proportion of SiO2 in the system increases. The dissolution of gold in aqueous SiO2 solutions can be described by the reaction: $$\begin{gathered} Au + 1/4O_2 + H_4 SiO_4^0 \rightleftharpoons AuH_3 SiO_4^0 + 1/2H_2 O \hfill \\ log K_{(200^\circ C)} = 6.23 \hfill \\ \end{gathered} $$ which indicates that SiO2 precipitation is an effective mechanism governing gold deposition, and thus explains the close association of silicification and gold mineralization.  相似文献   

12.
Reversed phase equilibrium experiments in the system (Ca, Mg, Fe)2SiO4 provide four tielines at P?1 bar and 1 kbar and 800° C–1,100° C. These tielines have been used to model the solution properties of the olivine quadrilateral following the methods described by Davidson et al. (1981) for quadrilateral clinopyroxenes. The discrepancy between the calculated phase relations and the experimentally determined tielines is within the uncertainty of the experiments. The solution properties of quadrilateral olivines can be described by a non-convergent site-disorder model that allows for complete partitioning of Ca on the M2 site, highly disordered Fe-Mg cation distributions and limited miscibility between high-Ca and low-Ca olivines. The ternary data presented in this paper together with binary solution models for the joins Fo-Mo and Fa-Kst have been used to evaluate two solution parameters: $$\begin{gathered} F^0 \equiv 2(\mu _{{\rm M}o}^0 - \mu _{{\rm K}st}^0 ) + \mu _{Fa}^0 - \mu _{Fo}^0 = 12.660 (1.6) kJ, \hfill \\ \Delta G_*^0 \equiv \mu _{{\rm M}gFe}^0 + \mu _{FeMg}^0 - \mu _{Fo}^0 - \mu _{Fa}^0 = 7.030 (3.9) kJ. \hfill \\ \end{gathered} $$ . Ternary phase quilibrium data for olivines tightly constrain the value of F0, but not that for ΔG * 0 which describes nonideality in Fe-Mg mixing. From this analysis, we infer a function for the apparent standard state energy of Kst: $$\begin{gathered} \mu _{{\rm K}st}^0 = - 102.79 \pm 0.8 - (T - 298)(0.137026) \hfill \\ + (T - 298 - T1n(T/298))(0.155519) \hfill \\ + (T - 298)^2 (2.8242E - 05)/2 \hfill \\ + (T - 298)^2 (2.9665E + 03)/(2T(298)^2 ) kJ \hfill \\ \end{gathered} $$ where T is in Kelvins and the 298 K value is relative to oxides.  相似文献   

13.
Experimental exchanges between plagioclases (synthesized from gels) and aqueous solutions (0.5N–8N) were carried out according to the reaction $$\begin{gathered} 2NaA1Si_3 O_8 + CaC1_2 \hfill \\ \leftrightarrow CaA1_2 Si_2 O_8 + 4SiO_2 + 2NaC1. \hfill \\ \end{gathered}$$ Distribution coefficients defined by $$K_D = \frac{{X_{An} }}{{(X_{Ab} )^2 }}\frac{{(X_{NaC1} )^2 }}{{X_{CaC1_2 } }}$$ were determined at 700° C and 1 kbar. From previous studies it is known that variations in the concentration of the aqueous solutions have no influence upon K D if the fluid is a single phase. In this study, variation of K D with the concentration of the solutions is interpreted as the result of fluid unmixing to vapour and brine phases. This implies boiling of CaCl2-NaCl-H2O fluids analogous to that known for the system NaCl-H2O. Experimental data permit calculation of the compositions of vapours and estimation of those of the brines for fluids in which Ca/Na<0.5. Boiling has an effect upon the exchange between feldspars and solutions (metasomatism) and must be considered when determining the activity coefficients.  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.
Precious metals accompany all types of epithermal deposits. In general, the largest of these deposits occur in intrusive or extrusive rocks of alkaline or calc-alkaline affinity. The Apigania Bay vein system and Au–Ag mineralization is hosted in Mesozoic marbles and schists, and is composed primarily of five nearly parallel, high-angle quartz veins that extend for at least 200 m. Gold–silver mineralization, in association with more than thirty ore and vein minerals, is developed in three stages and occurs at the contact of marbles and schists. Zones of epidote–chlorite–calcite and sericite–albite alteration are associated with precious metal-bearing milky and clear quartz veins. Fluid inclusion studies suggest that hydrothermal mineralization was deposited under hydrostatic pressures of ~100 bars, at temperature of 120–235°C, from low to moderate, calcium-bearing, saline fluids of 0.2 to 6.8 equiv. wt.% NaCl. Calculated isotope compositions (δ18O?=??4.7‰ to 1.7‰ and δD?=??120‰ to ?80‰) for waters in equilibrium with milky and clear quartz are consistent with mixing with dilute, low temperature meteoric ore fluids. Calculated δ 13CCO2 (0.6‰ to 1.1‰) and δ 34SH2S (?7.3 to ?0.3‰) compositions of the ore fluids indicate exchange, in an open system, with a metasedimentary source. Gold and silver deposition was associated with degassing of hydrogen due to intense uplift of the mineralizing area. The physicochemical conditions of mineralization stages I to III range between 200°C and 150°C, $f_{{\text{S}}_2 } = 10^{ - 18.1} $ to 10?16.8, $f_{{\text{O}}_2 } = 10^{ - 44.0} $ to 10?41.5, pH?=?6.9 to7.6, $f_{{\text{H}}_{\text{2}} {\text{S}}} = 10^{ - 3.4} $ to 10?2.6 and $a_{{\text{H}}_{\text{2}} {\text{S}}} = 10^{ - 2.7} $ to 10?2.6. Apigania Bay could be possibly considered the latest evolutional phase of Tinos hydrothermal system.  相似文献   

16.
The standard enthalpies of formation of FeS (troilite), FeS2 (pyrite), Co0.9342S, Co3S4 (linnaeite), Co9S8 (cobalt pentlandite), CoS2 (cattierite), CuS (covellite), and Cu2S (chalcocite) have been determined by high temperature direct reaction calorimetry at temperatures between 700 K and 1021 K. The following results are reported: $$\Delta {\rm H}_{f,FeS}^{tr} = - 102.59 \pm 0.20kJ mol^{ - 1} ,$$ $$\Delta {\rm H}_{f,FeS}^{py} = - 171.64 \pm 0.93kJ mol^{ - 1} ,$$ $$\Delta {\rm H}_{f,Co_{0.934} S} = - 99.42 \pm 1.52kJ mol^{ - 1} ,$$ $$\Delta {\rm H}_{f,Co_9 S_8 }^{ptl} = - 885.66 \pm 16.83kJ mol^{ - 1} ,$$ $$\Delta {\rm H}_{f,Co_3 S_4 }^{In} = - 347.47 \pm 7.27kJ mol^{ - 1} ,$$ $$\Delta {\rm H}_{f,CoS_2 }^{ct} = - 150.94 \pm 4.85kJ mol^{ - 1} ,$$ $$\Delta {\rm H}_{f,Cu_2 S}^{cc} = - 80.21 \pm 1.51kJ mol^{ - 1} ,$$ and $$\Delta {\rm H}_{f,CuS}^{cv} = - 53.14 \pm 2.28kJ mol^{ - 1} ,$$ The enthalpy of formation of CuFeS2 (chalcopyrite) from (CuS+FeS) and from (Cu+FeS2) was determined by solution calorimetry in a liquid Ni0.60S0.40 melt at 1100 K. The results of these measurements were combined with the standard enthalpies of formation of CuS, FeS, and FeS2, to calculate the standard enthalpy of formation of CuFeS2. We found \(\Delta {\rm H}_{f,CuFeS_2 }^{ccp} = - 194.93 \pm 4.84kJ mol^{ - 1}\) . Our results are compared with earlier data given in the literature; generally the agreement is good and our values agree with previous estimates within the uncertainties present in both.  相似文献   

17.
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   

18.
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   

19.
Data systematization using the constraints from the equation $$Cp = Cv + \alpha _P {}^2V_T K_T T$$ where C p, C v, α p, K T and V are respectively heat capacity at constant pressure, heat capacity at constant volume, isobaric thermal expansion, isothermal bulk modulus and molar volume, has been performed for tungsten and MgO. The data are $$K_T (W) = 1E - 5/(3.1575E - 12 + 1.6E - 16T + 3.1E - 20T^2 )$$ $$\alpha _P (W) = 9.386E - 6 + 5.51E - 9T$$ $$C_P (W) = 24.1 + 3.872E - 3T - 12.42E - 7T^2 + 63.96E - 11T^3 - 89000T^{ - 2} $$ $$K_T (MgO) = 1/(0.59506E - 6 + 0.82334E - 10T + 0.32639E - 13T^2 + 0.10179E - 17T^3 $$ $$\alpha _P (MgO) = 0.3754E - 4 + 0.7907E - 8T - 0.7836/T^2 + 0.9148/T^3 $$ $$C_P (MgO) = 43.65 + 0.54303E - 2T - 0.16692E7T^{ - 2} + 0.32903E4T^{ - 1} - 5.34791E - 8T^2 $$ For the calculation of pressure-volume-temperature relation, a high temperature form of the Birch-Murnaghan equation is proposed $$P = 3K_T (1 + 2f)^{5/2} (1 + 2\xi f)$$ Where $$K_T = 1/(b_0 + b_1 T + b_2 T^2 + b_3 T^3 )$$ $$f = (1/2)\{ [V(1,T)/V(P,T)]^{2/3} - 1\} $$ $$\xi = ({3 \mathord{\left/ {\vphantom {3 4}} \right. \kern-\nulldelimiterspace} 4})[K'_0 + K'_1 \ln ({T \mathord{\left/ {\vphantom {T {300}}} \right. \kern-\nulldelimiterspace} {300}}) - 4]$$ where in turn $$V(1,T) = V_0 [\exp (\int\limits_{300}^T {\alpha dT)]} $$ . The temperature dependence of the pressure derivative of the bulk modulus (K′1) is estimated by using the shock-wave data. For tungsten the data are K′0 = 3.5434, K′1 = 0.032; for MgO K′0 = 4.17 and K′1 = 0.1667. For calculating the Gibbs free energy of a solid at high pressure and at temperatures beyond that of melting at 1 atmosphere, it is necessary to define a high-temperature reference state for the fictive solid.  相似文献   

20.
The isotopic composition of calcite from travertine deposits of the Tokhana-Verkhnii hot spring in the Elbrus area shows broad variations in δ13C and δ18O (from +3.8 to +16.3‰ and from +24.6 to +28.1‰, respectively). The δ13C and δ18O values increase toward the sole of the travertine dome. The isotopically heaviest carbonates (δ13C of up to +16.3‰) were found near the bottom of the dome and composed ancient travertine, which are now not washed by mineral water. The scatter of the δ13C values of the fresh sample is slightly narrower: from +3.8 to +10‰. Calculations indicate that all carbonates of the Tokhana dome were not in equilibrium with spontaneous carbon dioxide released by the spring (\(\delta ^{13} C_{CO_2 } \) = ?8‰). To explain the generation of isotopically heavy travertine, a physicochemical model was developed for precipitation of Ca carbonates during the gradual degassing of the mineral water. The character of variations in the calculated δ13C values (from +5.5 to +13‰) is in good agreement with the tendency in the variations of the δ13C in the carbonate samples. The calculated and measured pH values are also consistent. Our results demonstrate that the isotopic composition of large travertine masses can be heterogeneous, and this should be taken into account during paleoclimatic and paleohydrogeological reconstruction.  相似文献   

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