Bimetasomatic zoning in the CaO-MgO-SiO2-H2O-CO2 system: Experiments with the use of natural rock samples

Experiments reproducing the development of bimetasomatic zoning in the CaO-MgO-SiO₂-H₂O-CO₂ system were conducted at elevated P-T parameters with the use of samples of naturally occurring quartzdolomite and calcite-serpentinite rocks. In order to maintain mass transfer exclusively via the diffusion-controlled mechanism, we used the method of the ensured compaction of the cylindrical sample surface with a thin-walled gold tube. In the course of the experiments, a single diopside zone ～2.5 × 10^?5 m thick was obtained at the quartz-dolomite interface at T = 600°C, $P_{H_2 O + CO_2 } $ = 200 MPa, and $X_{CO_2 } $ = 0.5 for 25–40 days and a succession of metasomatic zones at T = 750°C, $P_{H_2 O + CO_2 } $ = 300 MPa, and $X_{CO_2 } $ = 0.4 for 48 days. The metasomatic zones were as follows (listed in order from quartz to dolomite): wollastonite ‖ diopside ‖ tremolite ‖ calcite + forsterite; with the average width of the diopside zone equal to ～1.3 × 10^?5 m and the analogous part of the wollastonite zone equal to ～2.6 × 10^?5 m. Two zones (listed in order from calcite to serpentine) diopside and diopside-forsterite (the average widths of these zones were ～6 × 10^?4 and ～8 × 10^?4 m, respectively) were determined to develop at contact between serpentine and calcite during experiments that lasted 124 days at T = 500°C, $P_{H_2 O + CO_2 } $ = 200 MPa, and $X_{CO_2 } $ = 0.2–0.4. In the former and latter situations, the growth rate of the zoning ranged between 3.1 × 10^?12 and 1.2 × 10^?11 m/s and between 5.6 × 10^?11 and 7.5 × 10^?11 m/s, respectively. The higher growth rate in the latter case can be explained by the higher water mole fraction in the fluid, with this water released during serpentinite decomposition in the experiments. The development of the only diopside zone in the experiments modeling the interaction of quartz and dolomite at T = 600–650°C and $P_{H_2 O + CO_2 } $ = 200 MPa is in conflict with theoretical considerations underlain by the Korzhinskii-Fisher-Joesten model. The interaction of quartz and dolomite in the CaO-MgO-SiO₂-CO₂-H₂O system at the P-T- $X_{CO_2 } $ parameters specified above should be attended by the origin of a number of reaction zones consisting of various proportions of talc, forsterite, tremolite, diopside, and calcite. The saturation of the fluid with respect to these minerals was likely not reached, and this resulted in the degeneration of the respective stability fields in the succession of zones. Conceivably, this was related to the insufficient rates of quartz and dolomite dissolution and the relatively low diffusion rates of the dissolved species in the low-permeable medium. In the experiments with interacting calcite and serpentine, the zoning calcite ‖ diopside ‖ diopside + forsterite ‖ serpentine developed in its complete form, in agreement with the theory. Equilibrium was likely achieved in these experiments due to the higher diffusion coefficients.     

Crystal chemistry of perovskite-type compounds in the tausonite-loparite series, (Sr1−2 x Na x La x )TiO3

Perovskite-type compounds in the series tausonite-loparite, (Sr_{1?2 x} Na _x La _x )TiO₃, were synthesized by solid-state reaction (final heating at 1200–1300?°C), and studied using “conventional” and synchrotron X-ray powder diffractometry. The structures of intermediate compositions were determined using the Rietveld profile refinement method. In the compositional range 0?≤x?≤ 0.1, the series comprises perovskites characterized by an undistorted cubic structure (space group Pm3¯m, a?≈ 3.905–3.902?Å, Z?=?1). Intermediate compounds in the range 0.15?≤?x?≤?0.35 crystallize with tetragonal symmetry (I4/mcm, a?≈? , c?≈? , Z?=?4) derived from the cubic aristotype by antiphase rotation of the TiO₆ octahedra about a fourfold axis. The angle of rotation estimated from the positional parameters of oxygen atoms ranges from 2.5(7)° to 5.5(4)°. The cubic-to-tetragonal transition arises from substitution of Sr²⁺ by the comparatively smaller Na¹⁺ and La³⁺ cations. A further transition from the tetragonal to rhombohedral symmetry (R3¯c, a?≈? , c?≈?2 , Z?=?6) occurs between x?=?0.35 and 0.40, and apparently does not involve formation of perovskite with an intermediate two-tilt structure (Imma). The rhombohedral structure is characterized by a multicomponent octahedral tilt about a threefold axis ranging in magnitude from 6.5(2)° to 7.7(2)°. In the series (Sr_{1?2 x} Na _x La _x )TiO₃, the unit-cell dimensions decrease, and the degree of structural distortion increases with x.     

Thermodynamics of feldspathoid solutions

We have developed models for the thermody-namic properties of nephelines, kalsilites, and leucites in the simple system NaAlSiO₄?KAlSiO₄?Ca_0.5AlSiO₄?SiO₂?H₂O that are consistent with all known constraints on subsolidus equilibria and thermodynamic properties, and have integrated them into the existing MELTS software package. The model for nepheline is formulated for the simplifying assumptions that (1) a molecular mixing-type approximation describes changes in the configurational entropy associated with the coupled exchange substitutions □Si?NaAl and □Ca? Na₂ and that (2) Na⁺ and K⁺ display long–range non-convergent ordering between a large cation and the three small cation sites in the Na₄Al₄Si₄O₁₆ formula unit. Notable features of the model include the prediction that the mineral tetrakalsilite (“panunzite”, sensu stricto) results from anti-ordering of Na and K between the large cation and the three small cation sites in the nepheline structure at high temperatures, an average dT/dP slope of about 55^°/kbar for the reaction over the temperature and pressure ranges 800–1050 ^°C and 500–5000 bars, roughly symmetric (i.e. quadratic) solution behavior of the K–Na substitution along joins between fully ordered components in nepheline, and large positive Gibbs energies for the nepheline reciprocal reactions and and for the leucite reciprocal reaction      

Experimental investigation of the mechanism of the reaction: 1 tremolite+11 dolomite ⇌ 8 forsterite+13 calcite+9 CO2+1H2O

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.     

Interdiffusion of NaSi—CaAl in peristerite

The ‘average’ interdiffusion coefficient ( \(\bar D\) ) for NaSi—CaAl exchange in plagioclase for the interval from An₀ to An₂₆ was estimated from experimentally determined homogenization times for peristerite exsolution lamellae. The average spacing between adjacent (unlike) lamellae is 554±77 Å. Dry heating in air at 1,100°C for 98 days produced no change in the exsolution microstructure; thus \(\bar D\) (dry)<10^?17 cm²/s. This limit is consistent with the recently reported ‘average’ \(\bar D\) (dry) values for the Huttenlocher interval (An_70–90) at this temperature. At 1.5 GPa with about 0.2 weight percent water added the ‘average’ diffusion coefficient from 1,100°C to 900°C is given by: \(\bar D\) (wet)=18 _?15 ⁺¹⁰⁸ (cm²/s) exp (?97±5 (kcal/mol)/RT), where R is the gas constant, and T is °K. This \(\bar D\) (wet) at 1,100°C is more than three orders of magnitude greater than \(\bar D\) (dry) for Na- and Ca-rich plagioclases.     

Experimental investigation of the reaction forsterite + H2O ⇌ serpentine + brucite

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.     

EPR study of chromium-doped forsterite crystals: Cr3+(M1) with associated trivalent ions Al3+ and Sc3+

Electron paramagnetic resonance (EPR) study of single crystals of forsterite co-doped with chromium and scandium has revealed, apart from the known paramagnetic centers Cr³⁺(M1) and Cr³⁺(M1)– $ V_{{{\text{Mg}}^{2 + } }} $ (M2) (Ryabov in Phys Chem Miner 38:177–184, 2011), a new center Cr³⁺(M1)– $ V_{{{\text{Mg}}^{2 + } }} $ (M2)–Sc³⁺ formed by a Cr³⁺ ion substituting for Mg²⁺ at the M1 structural position with a nearest-neighbor Mg²⁺ vacancy at the M2 position and a Sc³⁺ ion presumably at the nearest-neighbor M1 position. For this center, the conventional zero-field splitting parameters D and E and the principal g values have been determined as follows: D?=?33,172(29) MHz, E?=?8,482(13) MHz, g?=?[1.9808(2), 1.9778(2), 1.9739(2)]. The center has been compared with the known ion pair Cr³⁺(M1)–Al³⁺ (Bershov et al. in Phys Chem Miner 9:95–101, 1983), for which the refined EPR data have been obtained. Based on these data, the known sharp M1″ line at 13,967?cm^?1 (with the splitting of 1.8?cm^?1), observed in low-temperature luminescence spectra of chromium-doped forsterite crystals (Glynn et al. in J Lumin 48, 49:541–544, 1991), has been ascribed to the Cr³⁺(M1)–Al³⁺ center. It has been found that the concentration of the new center increases from 0 up to 4.4?×?10¹⁵?mg^?1, whereas that of the Cr³⁺(M1) and Cr³⁺(M1)– $ V_{{{\text{Mg}}^{2 + } }} $ (M2) centers quickly decreases from 7.4?×?10¹⁵?mg^?1 down to 3?×?10¹⁵?mg^?1 and from 2.7?×?10¹⁵?mg^?1 down to 0.5?×?10¹⁵?mg^?1, i.e., by a factor of 2.5 and 5.4, respectively, with an increase of the Sc content from 0 up to 0.22 wt?% (at the same Cr content 0.25 wt?%) in the melt. When the Sc content exceeds that of Cr, the concentration of the new center decreases most likely due to the formation of the Sc³⁺(M1)– $ V_{{{\text{Mg}}^{2 + } }} $ (M2)–Sc³⁺ complex instead of the Cr³⁺(M1)– $ V_{{{\text{Mg}}^{2 + } }} $ (M2)–Sc³⁺ center. The formation of such ordered neutral complex is in agreement with the experimental results, concerning the incorporation of Sc into olivine, recently obtained by Grant and Wood (Geochim Cosmochim Acta 74:2412–2428, 2010).     

Experimental investigation of the metamorphism of siliceous dolomites

For the reaction: 1 diopside+3 dolomite ?2 forsterite+4 calcite+2 CO₂ (14) the following P _total?T-brackets have been determined experimentally in the presence of a gasphase consisting of 90 mole%CO₂ and 10 mole%H₂O∶1 kb, 544°±20° C; 3kb, 638°±15° C; 5kb, 708°±10° C; 10kb, 861°±10° C. The determination was carried out with well defined synthetic minerals in the starting mixture. The MgCO₃-contents of the magnesian calcites formed by the reaction in equilibrium with dolomite agree very well with the calcite-dolomite miscibility gap, which can be recalculated from the activities and the activity coefficients of MgCO₃ as given by Gordon and Greenwood (1970). The equilibrium constant K _14b was calculated with respect to the reference pressure P ₀=1 bar using the experimentally determined \(P_{total} TX_{CO_2 }\) brackets, the activities of MgCO₃ and CaCO₃ (Gordon and Greenwood 1970; Skippen 1974) and the fugacities of CO₂ Holloway (1977) considering the correction of Flowers (1979). Results are plotted as function of the absolute reciprocal temperature in Fig. 1. For the temperature range of 530° to 750° C the following linear expression can be given for the natural logarithm of K_14b: (g) $$[ln K_{14b} ]_T^P = - \frac{{18064.43}}{{T\left( {^\circ K} \right)}} + 38.58 + \frac{{0.308(P - 1 bar)}}{{T\left( {^\circ K} \right)}}$$ where P is the total pressure in bars and T the temperature in degrees Kelvin. Combining Equation (g) with the activities of MgCO₃ and CaCO₃ gives the equilibrium fugacity \(f_{CO_2 }\) : (i) $$[ln f_{CO_2 } ]_T^P = - \frac{{11635.44}}{{T\left( {^\circ K} \right)}} + 21.09 + \frac{{0.154(P - 1 bar)}}{{T\left( {^\circ K} \right)}}$$ Equation (i) and the fugacities of CO₂ permit to calculate the equilibrium data in terms of \(P_{CO_2 }\) and T (see Fig. 3) or P _total, T and \(X_{CO_2 }\) (see Fig. 5). Combining the \(P_{total} TX_{CO_2 }\) equilibrium data of the above reaction with those of the previously investigated reaction (Metz 1976): 1 tremolite+11 dolomite ?8 forsterite+13 calcite+9 CO₂+1 H₂O yields the stability conditions of the four-mineral assemblage: diopside+calcian dolomite+forsterite +magnesian calcite and the stability conditions of the five-mineral assemblage: tremolite+calcian dolomite+forsterite +magnesian calcite+diopside both shown in Fig. 6. Since these assemblages are by no means rare in metamorphic siliceous dolomites (Trommsdorff 1972; Suzuki 1977; Puhan 1979) the data of Fig. 6 can be used to determine the pressure of metamorphism and to estimate the composition of the CO₂-H₂O fluid provided the temperature of the metamorphic event was determined using the calcite-dolomite geothermometer.     

Experimental Na-K distribution between amphiboles and aqueous chloride solutions, and a mixing model along the richterite – K-richterite join

The cation exchange equilibrium has been investigated by hydrothermal experiments at 700 and 800°C at 200 MPa. To avoid equilibration problems of conventional exchange experiments, we synthesized amphiboles with an excess fluid allowing exchange between solid and fluid during the experiment. The exchangeable cations Na and K were provided as excess 1 to 2n chloridic solution. These exchange syntheses can be described by the reaction equation with ^(aq) for hydroxides and chlorides in aqueous solutions and _{( s )} and _{( p )}?=?start and product fluid. The amphiboles grew in presence of the exchange fluid and adjusted their stoichiometry in equilibrium with the fluid phase. The solid products consist of more than 99% amphibole (Na,K-richterite_ss) with traces of diopside and quartz. The amphiboles are up to 1?mm long and often ≈ 40 μm thick. Detailed EMP- and HRTEM-observations show that they are chemically homogeneous and structurally wellordered. The experimental results give consistent phase relations in the reciprocal ternary system Na-richterite–K-richterite–NaCl–KCl. We analysed the product fluid with AAS- and ICP-methods. The Na-K distribution coefficients between fluid and amphiboles of the richterite–K-richterite join are close to unity at 700°C and 800°C at 200 MPa. Small systematic deviations are explained by a symmetric solution model for the A-position of the amphiboles. Using ideal mixing for H₂O-NaCl-KCl fluids, a mixing model for the system richterite–K-richterite is presented. We suggest that the composition of richterite solid solutions can be used as a sensor for NaCl/KCl-ratios in metamorphic fluids.     

Gehlenite stability in the system CaO-Al2O3-SiO2-H2O-CO2

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.     

The ferric-ferrous ratio of natural silicate liquids equilibrated in air

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.     

The partial molar volume, thermal expansivity, and compressibility of H2O in NaAlSi3O8 liquid: new measurements and an internally consistent model

The molar volumes of 19 hydrous albitic liquids (1.9 to 6.1 wt% H₂O^total) were determined at one bar and 505–765 K. These volume data were derived from density measurements on hydrous glasses at 298 K, followed by measurements of the thermal expansion of each glass from 298 K to its respective glass transition temperature. The technique exploits the fact that the volume of a glass is equal to that of the corresponding liquid at the limiting fictive temperature (T _f′), and that T _f′ can be approximated as the temperature near the onset of the rapid increase in thermal expansion that occurs in the glass transition interval. The volume data of this study were combined with available volume data for anhydrous, Na₂O-Al₂O₃-SiO₂ liquids to derive the partial molar volume (±1) of the H₂O component in an albitic melt at ∼565 K and one bar. To extend the determination of to higher temperatures and pressures, the molar volumes of the hydrous albitic liquids determined in this study were combined with those measured by previous authors at 1023–1223 K and 480–840 MPa, leading to the following fitted values (±1) at 1673 K and one bar: (±0.46)×10⁻³ cm⁻³/mol-K, and dVˉ ^H ₂ ^{O total} /dP=−3.82 (±0.36)×10⁻⁴ cm³/mol-bar. The measured molar volumes of this study and those of previous authors can be recovered with a standard deviation of 0.5%, which is within the respective experimental errors. There is a significant difference between the values for derived in this study as a function of temperature and pressure and those obtained from an existing polynomial, primarily caused by the previous absence of accurate density measurements on anhydrous silicate liquids. The coefficients of thermal expansion (=4.72×10⁻⁴/K) and isothermal compressibility ( _T =1.66×10⁻⁵/bar) for the H₂O component at 1273 K and 100 MPa, indicate that H₂O^total is the single most expansive and compressible component in silicate liquids. For example, at 1473 K and 70 MPa (conditions of a mid-ocean ridge crustal magma chamber), the presence of just 0.4 wt% H₂O will decrease the density of a basaltic liquid by more than one percent. An equivalent decrease in melt density could be achieved by increasing the temperature by 175 degrees or the decreasing pressure by 230 MPa. Therefore, even minor quantities of dissolved water will have a marked effect on the dynamic properties of silicate liquids in the crustal environment. Received: 20 August 1996 / Accepted: 15 March 1997     

Experimental study of subsolidus phase relations and mixing properties of pyroxene and plagioclase in the system Na2O-CaO-Al2O3-SiO2

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.     

Experimentelle Untersuchung der Reaktion Dolomit + Kalifeldspat + H2O ⇌ Phlogopit + Calcit + CO2

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.     

On the nonstoichiometry and point defects of olivine

This paper presents the point-defect thermodynamics for fayalite and olivine solid solutions (Fe _x Mg_1?x )₂SiO₄. By means of thermogravimetry, the metal-to-oxygen ratio of these silicates has been determined as a function of oxygen potential, compositionx and temperature. Experiments were performed in the range of 1,000° C≦T≦1,280° C and 0.2≦x≦1.0. It is found that V _Me ^″ , Fe _Me ^· and the associate {Fe′ _Si ^′ Fe _Me ^· } are the majority defects. With this knowledge it is possible to calculate the nonstoichiometry at given temperature as a function of \(p_{O_2 } \) and \(a_{SiO_2 } \) . The cation vacancy concentration shows a \(p_{O_2 }^{1/5} \) -dependence (forx≧0.2) and increases at givenT and \(p_{O_2 } \) almost exponentially with compositionx. In the composition range studied here, the silicates show an oxygen excess, and FeO is more soluble in the olivine than SiO₂.     

Mg-ilmenite megacrysts from the Arkhangelsk kimberlites, Russia: Genesis and interaction with kimberlite melt and postkimberlite fluid

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.     

Later stages of evolution of an epithermal system: Au–Ag mineralizations at Apigania Bay, Tinos Island, Cyclades, Hellas, Greece

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 ¹³C_CO2 (0.6‰ to 1.1‰) and δ ³⁴S_H2S (?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.     

Experimental delineation of the “e” ⇌ I\bar 1 and “e” ⇌ C\bar 1 transformations in intermediate plagioclase feldspars

Approximately 125 hydrothermal annealing experiments have been carried out in an attempt to bracket the stability fields of different ordered structures within the plagioclase feldspar solid solution. Natural crystals were used for the experiments and were subjected to temperatures of ～650°C to ～1,000°C for times of up to 370 days at \(P_{{\text{H}}_{\text{2}} {\text{O}}} \) =600 bars, or \(P_{{\text{H}}_{\text{2}} {\text{O}}} \) =1,200 bars. The structural states of both parent and product materials were characterised by electron diffraction, with special attention being paid to the nature of type e and type b reflections (at h+k=(2n+1), l=(2n+1) positions). Structural changes of the type C \(\bar 1\) → I \(\bar 1\) , C \(\bar 1\) → “e” structure, I \(\bar 1\) → “e” and “e” structure → I \(\bar 1\) have been followed. There are marked differences between the ordering behaviour of crystals with compositions on either side of the C \(\bar 1\) ? I \(\bar 1\) transition line. In the composition range ～ An₅₀ to ～ An₇₀ the e structure appears to have a true field of stability relative to I \(\bar 1\) ordering, and a transformation of the type I \(\bar 1\) ? e has been reversed. It is suggested that the e structure is the more stable ordered state at temperatures of ～ 800°C and below. For compositions more albite-rich than ～ An₅₀ the upper temperature limit for long range e ordering is lower than ～ 750°C, and there is no evidence for any I \(\bar 1\) ordering. The evidence for a true stability field for “e” plagioclase, which is also consistent with calorimetric data, necessitates reanalysis both of the ordering behaviour of plagioclase crystals in nature and of the equilibrium phase diagram for the albite-anorthite system. Igneous crystals with compositions of ～ An₆₅, for example, probably follow a sequence of structural states C \(\bar 1\) → I \(\bar 1\) → e during cooling. The peristerite, Bøggild and Huttenlocher miscibility gaps are clearly associated with breaks in the albite, e and I \(\bar 1\) ordering behaviour but their exact topologies will depend on the thermodynamic character of the order/disorder transformations.     

High-temperature 29Si MAS NMR spectroscopy of anorthite (CaAl2Si2O8) and its P\bar 1-I\bar 1structural phase transition

We present ²⁹Si MAS NMR data for a well-ordered natural anorthite, obtained in situ at temperatures of from 25 to 500° C, which follow the changes in the aluminosilicate framework through the P $\bar 1$ -I $\bar 1$ structural phase transition. Pairs of peaks due to sites offset by approximately 1/2 [111] converge through the P $\bar 1$ phase and only four peaks are present above about 241° C. The variation of the peak positions with temperature and correlations based on structural data for the P $\bar 1$ and I $\bar 1$ phases allow assignment of all the MAS-NMR peaks to crystallographic sites. A Landau-type analysis gives an expression that relates the separation of pairs of con verging peaks to the local order parameter for the P $\bar 1$ -I $\bar 1$ transition, from which we determine its temperature dependence. Data for the best-constrained set of peak positions give for the order parameter critical exponent β = 0.27±0.04, consistent with previous results indicating that the P $\bar 1$ -I $\bar 1$ transition in pure anorthite is tricritical. No significant change in the ²⁹Si spin-lattice relaxation rate occurs across the P $\bar 1$ -I $\bar 1$ transition.     

High-temperature deformation of forsterite single crystals doped with vanadium

Creep experiments have been performed on samples from a single crystal of vanadium-doped forsterite under controlled \(p_{{\text{O}}_2 } \) conditions to investigate the effects of the addition of substitutional defects in the tetrahedral lattice sites. The addition of vanadium causes marked changes in the flow behavior of the forsterite, with a net increase in the creep rate at high \(p_{{\text{O}}_2 } \) and a new \(p_{{\text{O}}_2 } \) -dependent flow regime at low \(p_{{\text{O}}_2 } \) conditions. These observations can be interpreted as resulting from changes in the majority defect species that maintain the charge neutrality within the crystal. A climb-controlled dislocation creep model for the high-temperature deformation of vanadium-doped forsterite is proposed in which either (i) movement of uncharged jogs is rate-limited by the diffusion of silicon via a vacancy mechanism or (ii) movement of positively charged jogs is rate-limited by diffusion of oxygen via a vacancy mechanism.