Energetics of hydration of cordierite and water barometry in cordierite-granulites

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 \(P_{{\text{H}}_{\text{2}} {\text{O}}} \) =P_total, for assemblages from South India, Scottish Caledonides, Daly Bay and Hara Lake areas are compatible with those derived from the garnetplagioclase-sillimanite-quartz geobarometer.     

The system Fe-Si-O: Oxygen buffer calibrations to 1,500K

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 G_f, H_f, and S_f 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.     

The combined effect of F and H2O on interdiffusion between peralkaline dacitic and rhyolitic melts

The effective binary diffusion coefficient (EBDC) of silicon has been measured during the interdiffusion of peralkaline, fluorine-bearing (1.3 wt% F), hydrous (3.3 and 6 wt% H₂O), dacitic and rhyolitic melts at 1.0 GPa and temperatures between 1100°C and 1400°C. From Boltzmann-Matano analysis of diffusion profiles the diffusivity of silicon at 68 wt% SiO₂ can be described by the following Arrhenius equations (with standard errors): $$\begin{gathered} {\text{with 1}}{\text{.3 wt\% F and 3}}{\text{.3\% H}}_{\text{2}} {\text{O:}} \hfill \\ {\text{D}}_{{\text{Si}}} = \begin{array}{*{20}c} { + {\text{3}}{\text{.59}}} \\ {{\text{3}}{\text{.66}} \times {\text{10}}^{ - {\text{9}}} } \\ { - {\text{1}}{\text{.86}}} \\ \end{array} {\text{exp}}\left( {{{ - {\text{86}}{\text{.1}} \pm {\text{8}}{\text{.9}}} \mathord{\left/ {\vphantom {{ - {\text{86}}{\text{.1}} \pm {\text{8}}{\text{.9}}} {{\text{RT}}}}} \right. \kern-\nulldelimiterspace} {{\text{RT}}}}} \right) \hfill \\ {\text{with 1}}{\text{.3 wt\% F and 6}}{\text{.0\% H}}_{\text{2}} {\text{O:}} \hfill \\ {\text{D}}_{{\text{Si}}} = \begin{array}{*{20}c} { + {\text{3}}{\text{.59}}} \\ {{\text{3}}{\text{.51}} \times {\text{10}}^{ - {\text{8}}} } \\ { - {\text{1}}{\text{.77}}} \\ \end{array} {\text{exp}}\left( {{{ - {\text{109}}{\text{.5}} \pm {\text{8}}{\text{.9}}} \mathord{\left/ {\vphantom {{ - {\text{109}}{\text{.5}} \pm {\text{8}}{\text{.9}}} {{\text{RT}}}}} \right. \kern-\nulldelimiterspace} {{\text{RT}}}}} \right) \hfill \\ \end{gathered} $$ where D is in m²s^?1 and activation energies are in kJ/mol. Diffusivities measured at 64 and 72 wt% SiO₂ are only slightly different from those at 68 wt% SiO₂ and frequently all measurements are within error of each other. Silicon, aluminum, iron, magnesium, and calcium EBDCs were also calculated from diffusion profiles by error function inversion techniques assuming constant diffusivity. With one exception, silicon EBDCs calculated by error function techniques are within error of Boltzmann-Matano EBDCs. Average diffusivities of Fe, Mg, and Ca were within a factor of 2.5 of silicon diffusivities whereas Al diffusivities were approximately half those of silicon. Alkalies diffused much more rapidly than silicon and non-alkalies, however their diffusivities were not quantitatively determined. Low activation energies for silicon EBDCs result in rapid diffusion at magmatic temperatures. Assuming that water and fluorine exert similar effects on melt viscosity at high temperatures, the viscosity can be calculated and used in the Eyring equation used to determine diffusivities, typically to within a factor of three of those measured in this study. This correlation between viscosity and diffusivity can be inverted to calculate viscosities of fluorine- and water-bearing granitic melts at magmatic temperatures; these viscosities are orders of magnitude below those of hydrous granitic melts and result in more rapid and effective separation of granitic magmas from partially molten source rocks. Comparison of Arrhenius parameters for diffusion measured in this study with Arrhenius parameters determined for diffusion in similar compositions at the same pressure demonstrates simple relationships between Arrhenius parameters, activation energy-E_a, kJ/mol, pre-exponential factor-D_o, m²s^?1, and the volatile, X=F or OH^?, to oxygen, O, ratio of the melt {(X/X+O)}: $$\begin{gathered} {\text{E}}a = - {\text{1533\{ }}{{\text{X}} \mathord{\left/ {\vphantom {{\text{X}} {\left( {{\text{X}} + {\text{O}}} \right)}}} \right. \kern-\nulldelimiterspace} {\left( {{\text{X}} + {\text{O}}} \right)}}{\text{\} }} + {\text{213}}{\text{.3}} \hfill \\ {\text{D}}_{\text{O}} = {\text{2}}{\text{.13}} \times {\text{10}}^{ - {\text{6}}} {\text{exp}}\left[ { - {\text{6}}{\text{.5\{ }}{{\text{X}} \mathord{\left/ {\vphantom {{\text{X}} {\left( {{\text{X}} + {\text{O}}} \right)}}} \right. \kern-\nulldelimiterspace} {\left( {{\text{X}} + {\text{O}}} \right)}}{\text{\} }}} \right] \hfill \\ \end{gathered} $$ These relationships can be used to estimate diffusion in various melts of dacitic to rhyolitic composition containing both fluorine and water. Calculations for the contamination of rhyolitic melts by dacitic enclaves at 800°C and 700°C provide evidence for the virtual inevitability of diffusive contamination in hydrous and fluorine-bearing magmas if they undergo magma mixing of any form.     

The reaction garnet+clinopyroxene+quartz =2 orthopyroxene+anorthite: A potential geobarometer for granulites

A mineralogic geobarometer based on the reaction garnet+clinopyroxene+quartz=2 orthopyroxene+anorthite is proposed. The geobarometric formulations for the Fe- and Mg- end member equilibria are $$\begin{gathered} P_{({\text{Fe}})} {\text{ }}({\text{bars}}){\text{ = 32}}{\text{.097 }}T{\text{ }} - {\text{ 26385 }} - {\text{ 22}}{\text{.79 (}}T - 848 - T1{\text{n(}}T/848{\text{))}} \hfill \\ {\text{ }} - (3.655 + 0.0138T){\text{ }}\left( {\frac{{{\text{(}}T - 848{\text{)}}^{\text{2}} }}{T}} \right) \hfill \\ {\text{ }} - {\text{(3}}{\text{.123) }}T1{\text{n }}\frac{{(a_{a{\text{n}}}^{{\text{Plag}}} )(a_{{\text{fs}}}^{{\text{P}}\ddot u{\text{x}}} )^2 }}{{(a_{{\text{alm}}}^{{\text{Gt}}} )(a_{{\text{hed}}}^{{\text{Opx}}} )}} \hfill \\ P_{({\text{Mg}})} {\text{ (bars) = 9}}{\text{.270 }}T + 4006 - 0.9305{\text{ }}(T - 848 - T1{\text{n (}}T/848{\text{)}}) \hfill \\ {\text{ }} - (1.1963{\text{ }} - {\text{ }}6.0128{\text{ x 10}}^{ - {\text{3}}} T)\left( {\frac{{(T - 848)^2 }}{T}} \right) \hfill \\ {\text{ }} - 3.489{\text{ }}T1{\text{n }}\frac{{(a_{an}^{{\text{Plag}}} ){\text{ }}(a_{{\text{ens}}}^{{\text{Opx}}} )}}{{{\text{(}}a_{{\text{pyr}}}^{{\text{Gt}}} {\text{) (}}a_{{\text{diop}}}^{{\text{Cpx}}} {\text{)}}}}. \hfill \\ \end{gathered}$$ The end member thermodynamic data have been taken from the data base of Helgeson et al. (1978) and Saxena and Erikson (1983). The activities of pyroxene components and anorthite in plagioclase have been modelled after Wood and Banno (1973) and Newton (1983) respectively. The activities of pyrope and almandine are calculated from the binary interaction parameters for garnet solid solutions proposed by Saxena and Erikson (1983). Pressures computed from these equations for fifty sets of published mineral data from several granulite areas are comparable with those obtained from dependable geobarometers. The pressure values determined from the Fe-end member equilibrium appear to be more reasonable than those from the Mg-end member reaction. It is likely that the difference in pressures computed from the Fe- and Mg-end members, ΔP ^*, have been caused by non-ideal mixing in the phases, especially in garnets.     

The H<Subscript>2</Subscript>O solubility of alkali basaltic melts: an experimental study

Experiments were conducted to determine the water solubility of alkali basalts from Etna, Stromboli and Vesuvius volcanoes, Italy. The basaltic melts were equilibrated at 1,200°C with pure water, under oxidized conditions, and at pressures ranging from 163 to 3,842 bars. Our results show that at pressures above 1 kbar, alkali basalts dissolve more water than typical mid-ocean ridge basalts (MORB). Combination of our data with those from previous studies allows the following simple empirical model for the water solubility of basalts of varying alkalinity and fO₂ to be derived: \textH ₂ \textO( \textwt% ) = \text H ₂ \textO_\textMORB ( \textwt% ) + ( 5.84 ×10 ^{- 5} *\textP - 2.29 ×10 ^{- 2} ) ×( \textNa₂ \textO + \textK₂ \textO )( \textwt% ) + 4.67 ×10 ^{- 2} ×\Updelta \textNNO - 2.29 ×10 ^{- 1} {\text{H}}_{ 2} {\text{O}}\left( {{\text{wt}}\% } \right) = {\text{ H}}_{ 2} {\text{O}}_{\text{MORB}} \left( {{\text{wt}}\% } \right) + \left( {5.84 \times 10^{ - 5} *{\text{P}} - 2.29 \times 10^{ - 2} } \right) \times \left( {{\text{Na}}_{2} {\text{O}} + {\text{K}}_{2} {\text{O}}} \right)\left( {{\text{wt}}\% } \right) + 4.67 \times 10^{ - 2} \times \Updelta {\text{NNO}} - 2.29 \times 10^{ - 1} where H₂O_MORB is the water solubility at the calculated P, using the model of Dixon et al. (1995). This equation reproduces the existing database on water solubilities in basaltic melts to within 5%. Interpretation of the speciation data in the context of the glass transition theory shows that water speciation in basalt melts is severely modified during quench. At magmatic temperatures, more than 90% of dissolved water forms hydroxyl groups at all water contents, whilst in natural or synthetic glasses, the amount of molecular water is much larger. A regular solution model with an explicit temperature dependence reproduces well-observed water species. Derivation of the partial molar volume of molecular water using standard thermodynamic considerations yields values close to previous findings if room temperature water species are used. When high temperature species proportions are used, a negative partial molar volume is obtained for molecular water. Calculation of the partial molar volume of total water using H₂O solubility data on basaltic melts at pressures above 1 kbar yields a value of 19 cm³/mol in reasonable agreement with estimates obtained from density measurements.     

(Ca,Mg)2Si2O6 clinopyroxenes: A solution model based on nonconvergent site-disorder

Experiments at high pressure and temperature indicate that excess Ca may be dissolved in diopside. If the (Ca, Mg)₂Si₂O₆ clinopyroxene solution extends to more Ca-rich compositions than CaMgSi₂O₆, macroscopic regular solution models cannot strictly be applied to this system. A nonconvergent site-disorder model, such as that proposed by Thompson (1969, 1970), may be more appropriate. We have modified Thompson's model to include asymmetric excess parameters and have used a linear least-squares technique to fit the available experimental data for Ca-Mg orthopyroxene-clinopyroxene equilibria and Fe-free pigeonite stability to this model. The model expressions for equilibrium conditions \(\mu _{{\text{Mg}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} }^{{\text{opx}}} = \mu _{{\text{Mg}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} }^{{\text{cpx}}} \) (reaction A) and \(\mu _{{\text{Ca}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} }^{{\text{opx}}} = \mu _{{\text{Ca}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} }^{{\text{cpx}}} \) (reaction B) are given by: 1 $$\begin{gathered} \Delta \mu _{\text{A}}^{\text{O}} = {\text{RT 1n}}\left[ {\frac{{(X_{{\text{Mg}}}^{{\text{opx}}} )^2 }}{{X_{{\text{Mg}}}^{{\text{M1}}} \cdot X_{{\text{Mg}}}^{{\text{M2}}} }}} \right] - \frac{1}{2}\{ W_{21} [2(X_{{\text{Ca}}}^{{\text{M2}}} )^3 - (X_{{\text{Ca}}}^{{\text{M2}}} ] \hfill \\ {\text{ + 2W}}_{{\text{22}}} [X_{{\text{Ca}}}^{{\text{M2}}} )^2 - (X_{{\text{Ca}}}^{{\text{M2}}} )^3 + \Delta {\text{G}}_{\text{*}}^{\text{0}} (X_{{\text{Ca}}}^{{\text{M1}}} \cdot X_{{\text{Ca}}}^{{\text{M2}}} )\} \hfill \\ {\text{ + W}}^{{\text{opx}}} (X_{{\text{Wo}}}^{{\text{opx}}} )^2 \hfill \\ \Delta \mu _{\text{B}}^{\text{O}} = {\text{RT 1n}}\left[ {\frac{{(X_{{\text{Ca}}}^{{\text{opx}}} )^2 }}{{X_{{\text{Ca}}}^{{\text{M1}}} \cdot X_{{\text{Ca}}}^{{\text{M2}}} }}} \right] - \frac{1}{2}\{ 2W_{21} [2(X_{{\text{Mg}}}^{{\text{M2}}} )^2 - (X_{{\text{Mg}}}^{{\text{M2}}} )^3 ] \hfill \\ {\text{ + W}}_{{\text{22}}} [2(X_{{\text{Mg}}}^{{\text{M2}}} )^3 - (X_{{\text{Mg}}}^{{\text{M2}}} )^2 + \Delta {\text{G}}_{\text{*}}^{\text{0}} (X_{{\text{Mg}}}^{{\text{M1}}} \cdot X_{{\text{Mg}}}^{{\text{M2}}} )\} \hfill \\ {\text{ + W}}^{{\text{opx}}} (X_{{\text{En}}}^{{\text{opx}}} )^2 \hfill \\ \hfill \\ \end{gathered} $$ where 1 $$\begin{gathered} \Delta \mu _{\text{A}}^{\text{O}} = 2.953 + 0.0602{\text{P}} - 0.00179{\text{T}} \hfill \\ \Delta \mu _{\text{B}}^{\text{O}} = 24.64 + 0.958{\text{P}} - (0.0286){\text{T}} \hfill \\ {\text{W}}_{{\text{21}}} = 47.12 + 0.273{\text{P}} \hfill \\ {\text{W}}_{{\text{22}}} = 66.11 + ( - 0.249){\text{P}} \hfill \\ {\text{W}}^{{\text{opx}}} = 40 \hfill \\ \Delta {\text{G}}_*^0 = 155{\text{ (all values are in kJ/gfw)}}{\text{.}} \hfill \\ \end{gathered} $$ . Site occupancies in clinopyroxene were determined from the internal equilibrium condition 1 $$\begin{gathered} \Delta G_{\text{E}}^{\text{O}} = - {\text{RT 1n}}\left[ {\frac{{X_{{\text{Ca}}}^{{\text{M1}}} \cdot X_{{\text{Mg}}}^{{\text{M2}}} }}{{X_{{\text{Ca}}}^{{\text{M2}}} \cdot X_{{\text{Mg}}}^{{\text{M1}}} }}} \right] + \tfrac{1}{2}[(2{\text{W}}_{{\text{21}}} - {\text{W}}_{{\text{22}}} )(2{\text{X}}_{{\text{Ca}}}^{{\text{M2}}} - 1) \hfill \\ {\text{ + }}\Delta G_*^0 (X_{{\text{Ca}}}^{{\text{M1}}} - X_{{\text{Ca}}}^{{\text{M2}}} ) + \tfrac{3}{2}(2{\text{W}}_{{\text{21}}} - {\text{W}}_{{\text{22}}} ) \hfill \\ {\text{ (1}} - 2X_{{\text{Ca}}}^{{\text{M1}}} )(X_{{\text{Ca}}}^{{\text{M1}}} + \tfrac{1}{2})] \hfill \\ \end{gathered} $$ where δG _E ⁰ =153+0.023T+1.2P. The predicted concentrations of Ca on the clinopyroxene Ml site are low enough to be compatible with crystallographic studies. Temperatures calculated from the model for coexisting ortho- and clinopyroxene pairs fit the experimental data to within 10° in most cases; the worst discrepancy is 30°. Phase relations for clinopyroxene, orthopyroxene and pigeonite are successfully described by this model at temperatures up to 1,600° C and pressures from 0.001 to 40 kbar. Predicted enthalpies of solution agree well with the calorimetric measurements of Newton et al. (1979). The nonconvergent site disorder model affords good approximations to both the free energy and enthalpy of clinopyroxenes, and, therefore, the configurational entropy as well. This approach may provide an example for Febearing pyroxenes in which cation site exchange has an even more profound effect on the thermodynamic properties.     

The effect of H<Subscript>2</Subscript>O on the olivine liquidus of basaltic melts: experiments and thermodynamic models

We designed and carried out experiments to investigate the effect of H₂O on the liquidus temperature of olivine-saturated primitive melts. The effect of H₂O was isolated from other influences by experimentally determining the liquidus temperatures of the same melt composition with various amounts of H₂O added. Experimental data indicate that the effect of H₂O does not depend on pressure or melt composition in the basaltic compositional range. The influence of H₂O on melting point lowering can be described as a polynomial function This expression can be used to account for the effect of H₂O on olivine-melt thermometers, and can be incorporated into fractionation models for primitive basalts. The non-linear effect of H₂O indicates that incorporation of H₂O in silicate melts is non-ideal, and involves interaction between H₂O and other melt components. The simple speciation approach that seems to account for the influence of H₂O in simple systems (albite-H₂O, diopside-H₂O) fails to describe the mixing behavior of H₂O in multi-component silicate melts. However, a non-ideal solution model that treats the effect of H₂O addition as a positive excess free energy can be fitted to describe the effect of melting point lowering.     

Zinc isotope fractionation under vaporization processes and in aqueous solutions

Equilibrium Zn isotope fractionation was investigated using first-principles quantum chemistry methods at the B3LYP/6-311G* level. The volume variable cluster model method was used to calculate isotope fractionation factors of sphalerite, smithsonite, calcite, anorthite, forsterite, and enstatite. The water-droplet method was used to calculate Zn isotope fractionation factors of Zn²⁺-bearing aqueous species; their reduced partition function ratio factors decreased in the order \(\left[ {{\text{Zn}}\left( {{\text{H}}_{2} {\text{O}}} \right)_{6} } \right]^{2 + } > \left[ {{\text{ZnCl}}\left( {{\text{H}}_{2} {\text{O}}} \right)_{5} } \right]^{ + } > \left[ {{\text{ZnCl}}_{2} \left( {{\text{H}}_{2} {\text{O}}} \right)_{4} } \right] > \left[ {{\text{ZnCl}}_{3} \left( {{\text{H}}_{2} {\text{O}}} \right)_{2} } \right]^{ - } > {\text{ZnCl}}_{4} ]^{2 - }\). Gaseous ZnCl₂ was also calculated for vaporization processes. Kinetic isotope fractionation of diffusional processes in a vacuum was directly calculated using formulas provided by Richter and co-workers. Our calculations show that in addition to the kinetic isotope effect of diffusional processes, equilibrium isotope fractionation also contributed nontrivially to observed Zn isotope fractionation of vaporization processes. The calculated net Zn isotope fractionation of vaporization processes was 7–7.5‰, with ZnCl₂ as the gaseous species. This matches experimental observations of the range of Zn isotope distribution of lunar samples. Therefore, vaporization processes may be the cause of the large distribution of Zn isotope signals found on the Moon. However, we cannot further distinguish the origin of such vaporization processes; it might be due either to igneous rock melting in meteorite bombardments or to a giant impact event. Furthermore, isotope fractionation between Zn-bearing aqueous species and minerals that we have provided helps explain Zn isotope data in the fields of ore deposits and petrology.     

The Hydrolysis of Al(III) in NaCl solutions: A Model for M(II), M(III), and M(IV) Ions

Understanding the identity and stability of the hydrolysis products of metals is required in order to predict their behavior in natural aquatic systems. Despite this need, the hydrolysis constants of many metals are only known over a limited range of temperature and ionic strengths. In this paper, we show that the hydrolysis constants of 31 metals [i.e. Mn(II), Cr(III), U(IV), Pu(IV)] are nearly linearly related to the values for Al(III) over a wide range of temperatures and ionic strengths. These linear correlations allow one to make reasonable estimates for the hydrolysis constants of +2, +3, and +4 metals from 0 to 300°C in dilute solutions and 0 to 100°C to 5 m in NaCl solutions. These correlations in pure water are related to the differences between the free energies of the free ion and complexes being almost equal $$ \Updelta {\text{G}}^\circ \left( {{\text{Al}}^{3 + } } \right) - \Updelta {\text{G}}^\circ \left( {{\text{Al}}\left( {\text{OH}} \right)_{j}^{{\left( {3 - j} \right)}} } \right) \cong \Updelta {\text{G}}^\circ \left( {{\text{M}}^{n + } } \right) - \Updelta {\text{G}}^\circ \left( {{\text{M}}\left( {\text{OH}} \right)_{j}^{{\left( {n - j} \right)}} } \right) $$ The correlation at higher temperatures is a result of a similar relationship between the enthalpies of the free ions and complexes $$ \Updelta {\text{H}}^\circ \left( {{\text{Al}}^{3 + } } \right) - \Updelta {\text{H}}^\circ \left( {{\text{Al}}\left( {\text{OH}} \right)_{j}^{3 - j} } \right) \cong \Updelta {\text{H}}^\circ \left( {{\text{M}}^{n + } } \right) - \Updelta {\text{H}}^\circ \left( {{\text{M}}\left( {\text{OH}} \right)_{j}^{n - j} } \right) $$ The correlations at higher ionic strengths are the result of the ratio of the activity coefficients for Al(III) being almost equal to that of the metal. $$ \gamma \left( {{\text{M}}^{n + } } \right)/\gamma \left( {{\text{M}}\left( {\text{OH}} \right)_{j}^{n - j} } \right) \cong \gamma \left( {{\text{Al}}^{3 + } } \right)/\gamma \left( {{\text{Al}}\left( {\text{OH}} \right)_{j}^{3 - j} } \right) $$ The results of this study should be useful in examining the speciation of metals as a function of pH in natural waters (e.g. hydrothermal fresh waters and NaCl brines).     

Aluminum partitioning between olivine and ultrabasic silicate liquid to 6 GPa

Trace element analyses of 1-atm and high-pressure experiments show that in komatiite and peridotite, the olivine (OL)/liquid (L) distribution coefficient for Al₂O₃ ( ) increases with pressure and temperature. Olivine in equilibrium with liquid accepts as much as 0.2 wt% Al₂O₃ in solution at 6 GPa. Convergence to equilibrium compositions at this high level is shown by cation diffusion of Al into synthetic forsterite crystals of low-Al contents in the presence of melt. Convergence to low-Al equilibrium compositions at lower P and T is shown by diffusion of Al out of synthetic forsterite with high initial Al content. Isobaric and isothermal experimental data subsets reveal that temperature and pressure variations both have real effects on . Variation in silicate melt composition has no detectable effect on within the limited range of experimentally investigated mixtures. Least-squares regression for 24 experiments, using komatiite and peridotite, performed at 1 atm to 6 GPa and 1300 to 1960°C, gives the best fit equation: Increase in with increasingly higher-pressure melting is consistent with incorporation of a spinel-like component of low molar volume into olivine, although other substitutions possibly involving more complex coupling cannot be ruled out. High P-T ultrabasic melting residues, if pristine, may be recognized by the high calculated from microprobe analyses of Al₂O₃ concentrations in residual olivines and estimated Al₂O₃ concentration in the last liquid removed. In general the low levels of Al in natural olivine from mantle xenoliths suggest that pristine residues are rarely recovered.     

Temperature coefficients of elastic constants of single crystal MgO between 80 and 1,300 K

Elastic constants of single crystal MgO have been measured by the rectangular parallelepiped resonance (RPR) method at temperatures between 80 and 1,300 K. Elastic constants C _ij (Mbar=10³ kbar) and their temperature coefficients (kbar/K) are: $$\begin{gathered} {\text{ }}C_{{\text{11}}} {\text{ }}C_{{\text{12}}} {\text{ }}C_{{\text{44}}} {\text{ }}K_s {\text{ }}C_s \hfill \\ C_{ij} {\text{ 300 K 2}}{\text{.966 0}}{\text{.959 1}}{\text{.562 1}}{\text{.628 1}}{\text{.004}} \hfill \\ \partial C_{ij} {\text{/}}\partial T{\text{100 K }} - {\text{0}}{\text{.259 0}}{\text{.013 }} - {\text{0}}{\text{.072 }} - {\text{0}}{\text{.078 }} - {\text{0}}{\text{.136}} \hfill \\ {\text{ 300K }} - {\text{0}}{\text{.596 0}}{\text{.068 }} - {\text{0}}{\text{.122 }} - {\text{0}}{\text{.153 }} - {\text{0}}{\text{.332}} \hfill \\ {\text{ 800 K }} - {\text{0}}{\text{.619 0}}{\text{.009 }} - {\text{0}}{\text{.152 }} - {\text{0}}{\text{.200 }} - {\text{0}}{\text{.314}} \hfill \\ {\text{ 1,300 K }} - {\text{0}}{\text{.598 0}}{\text{.036 }} - {\text{0}}{\text{.130 }} - {\text{0}}{\text{.223 }} - {\text{0}}{\text{.218}} \hfill \\ \end{gathered} $$ By combining the present results with the previous data on the thermal expansivity and specific heat, the thermodynamic properties of magnesium oxide are presented and discussed. The elastic parameters of MgO at very high temperatures in the earth's lower mantle are also clarified.     

Statistical mechanics of coupled solid solutions in the dilute limit

A statistical mechanical analysis of the limiting laws for coupled solid solutions shows that the random model, in which the configurational entropy is calculated as if atoms mix randomly on each crystallographic site, is correct as a first approximation. In coupled solid solutions, since atoms of different valence substitute on the same sites, significant short-range order which reduces the entropy can be expected. A first-order correction is rigorously obtained for the entropy in dilute binary short-range ordered coupled solid solutions: $$\bar S^{{\text{XS}}} {\text{/R = }}Q\left( {{\text{e}}^{--H_{\text{A}} /{\text{R}}T} \left( {\frac{{H_{\text{A}} }}{{{\text{R}}T}} + 1} \right) - 1} \right)N_2^a N_4^b ,$$ where Q is the number of positions an associated cation pair can assume per formula unit, H _A is the association energy per formula unit, and N ₂ ^a and N ₄ ^b are the site occupancy fractions for atoms 2 and 4 that are dilute on sites a and b. S ^XS is the configurational entropy minus the random model entropy. Aluminous pyroxenes on the joints diopside-jadeite and diopside-CaTs are examined as examples. A generalization for dilute multiple component solutions, including possible long-range ordering variations is given by: $$\frac{{\bar S^{{\text{XS}}} }}{{\text{R}}}{\text{ = }}\sum\limits_i {\sum\limits_j {\sum\limits_k {Q_i } } \left( {{\text{e}}^{--H_{\text{A}}^{j{\text{ }}k{\text{, }}i} /{\text{R}}T} \left( {\frac{{H_{\text{A}}^{j{\text{ }}k{\text{, }}i} }}{{{\text{R}}T}} + 1} \right) - 1} \right)N_j^l N_k^m ,} $$ where i labels each crystallographically distinct pair, j and k label atomic species, l and m label crystallographic sites, and the N's are site occupancy fractions for the solute atoms. A total association model is examined as well as the partial association and random models. Real solution behavior must lie between the total association model and the random model. Molecular models in which the ideal activity is proportional to a mole fraction, which in itself is not always unambiguously defined, do not lie in this range and furthermore have no physical justification.     

The carbon dioxide solubility in alkali basalts: an experimental study

Experiments were conducted to determine CO₂ solubilities in alkali basalts from Vesuvius, Etna and Stromboli volcanoes. The basaltic melts were equilibrated with nearly pure CO₂ at 1,200°C under oxidizing conditions and at pressures ranging from 269 to 2,060 bars. CO₂ solubility was determined by FTIR measurements. The results show that alkalis have a strong effect on the CO₂ solubility and confirm and refine the relationship between the compositional parameter Π devised by Dixon (Am Mineral 82:368–378, 1997) and the CO₂ solubility. A general thermodynamic model for CO₂ solubility in basaltic melts is defined for pressures up to 2 kbars. Based on the assumption that O²⁻ and CO₃²⁻ mix ideally, we have:

Physicochemical parameter charts for gases in fluid inclusions

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.     

Non-ideal mixing in the phlogopite-annite binary: constraints from experimental data on Mg−Fe partitioning and a reformulation of the biotite-garnet geothermometer

The existing experimental data [Ferry and Spear 1978; Perchuk and Lavrent'eva 1983] on Mg?Fe partitioning between garnet and biotite are disparate. The underlying assumption of ideal Mg?Fe exchange between the minerals has been examined on the basis of recently available thermochemical data. Using the updated mixing parameters for the pyrope-almandine asymmetric regular solution as inputs [Ganguly and Saxena 1984; Hackler and Wood 1984], thermodynamic analysis points to non-ideal mixing in the phlogopite-annite binary in the temperature range of 550°C–950°C. The non-ideality can be approximated by a temperature-independent, one constant Margules parameter. The retrieved values for enthalpy of mixing for Mg?Fe biotites and the standard state enthalpy and entropy changes of the exchange reaction were combined with existing thermochemical data on grossular-pyrope and grossular-almandine binaries to obtain geothermometric expressions for Mg?Fe fractionation between biotite and garnet. [T in K] $$\begin{gathered} {\text{T(HW) = [20286 + 0}}{\text{.0193P - \{ 2080(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{)}}^{\text{2}} {\text{ - 6350(X}}_{{\text{Fe}}}^{{\text{Gt}}} {\text{)}}^{\text{2}} \hfill \\ {\text{ - 13807(X}}_{{\text{Ca}}}^{{\text{Gt}}} {\text{)(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} {\text{) + 8540(X}}_{{\text{Fe}}}^{{\text{Gt}}} {\text{)(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{)(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} {\text{)}} \hfill \\ {\text{ + 4215(X}}_{{\text{Ca}}}^{{\text{Gt}}} {\text{)(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{ - X}}_{{\text{Fe}}}^{{\text{Gt}}} {\text{)\} + 4441}}{{{\text{(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} \mathord{\left/ {\vphantom {{{\text{(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} {{\text{[13}}{\text{.138}}}}} \right. \kern-\nulldelimiterspace} {{\text{[13}}{\text{.138}}}} \hfill \\ {\text{ + 8}}{\text{.3143 InK}}_{\text{D}} {\text{ + 6}}{\text{.276(X}}_{{\text{Ca}}}^{{\text{Gt}}} ){\text{(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} )] \hfill \\ {\text{T(GS) = [13538 + 0}}{\text{.0193P - \{ 837(X}}_{{\text{Mg}}}^{{\text{Gt}}} )^{\text{2}} {\text{ - 10460(X}}_{{\text{Fe}}}^{{\text{Gt}}} )^2 \hfill \\ {\text{ - 13807(X}}_{{\text{Ca}}}^{{\text{Gt}}} )(1{\text{ - X}}_{{\text{Mn}}}^{{\text{Gt}}} {\text{) + 19246(X}}_{{\text{Fe}}}^{{\text{Gt}}} ){\text{(X}}_{{\text{Mg}}}^{{\text{Gt}}} ){\text{(1 - X}}_{{\text{Mn}}}^{{\text{Gt}}} ) \hfill \\ {\text{ }}{{{\text{ + 5649(X}}_{{\text{Ca}}}^{{\text{Gt}}} ){\text{(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{ - X}}_{{\text{Fe}}}^{{\text{Gt}}} ){\text{\} + 7972(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} \mathord{\left/ {\vphantom {{{\text{ + 5649(X}}_{{\text{Ca}}}^{{\text{Gt}}} ){\text{(X}}_{{\text{Mg}}}^{{\text{Gt}}} {\text{ - X}}_{{\text{Fe}}}^{{\text{Gt}}} ){\text{\} + 7972(2X}}_{{\text{Mg}}}^{{\text{Bt}}} {\text{ - 1)]}}} {{\text{[6}}{\text{.778}}}}} \right. \kern-\nulldelimiterspace} {{\text{[6}}{\text{.778}}}} \hfill \\ {\text{ + 8}}{\text{.3143InK}}_{\text{D}} {\text{ + 6}}{\text{.276(X}}_{{\text{Ca}}}^{{\text{Gt}}} )(1{\text{ - X}}_{{\text{Mn}}}^{{\text{Gt}}} )] \hfill \\ \end{gathered} $$ The reformulated geothermometer is an improvement over existing biotite-garnet geothermometers because it reconciles the experimental data sets on Fe?Mg partitioning between the two phases and is based on updated activity-composition relationship in Fe?Mg?Ca garnet solid solutions.     

Raman spectra of gypsum under pressure

Raman sprectra of a gypsum crystal were made at pressures between 0.001 and 7 kbar using He gas as the pressure medium. \(\frac{{{\text{d}}v}}{{dP}}\) values for bands in the range 3,600–100 cm^?1 were obtained. Comparison of results with \(\frac{{{\text{d}}v}}{{{\text{d}}T}}\) from the literature for temperatures of 77 and 300° K. shows that the internal modes of the SO₄ units are more sensitive to pressure than to temperature. The effect is small. Coupled H₂O-SO₄ translational modes are greatly affected by both pressure and temperature while coupled Ca-SO₄ mode are less so. It was found that stretching vibrations of water molecules were affected differently under pressure. The band at 3,500 cm^?1 is more greatly displaced by pressure \(\left( {\frac{{{\text{d}}v}}{{{\text{d}}P}} = {\text{2}}{\text{.11cm}}^{{\text{ - 1}}} /{\text{kbar}}} \right)\) than the band at 3,400 cm^?1 \(\left( {\frac{{{\text{d}}v}}{{{\text{d}}P}} \simeq {\text{2}}{\text{.11cm}}^{{\text{ - 1}}} /{\text{kbar}}} \right)\) . Assuming two different hydrogen bond intensities for the water molecules, one can attribute this difference in behavior of stretching modes to and increase in hydrogen bonding of one of the hydrogens which is exterior to the double H₂O planes in the gypsum structure. The great variety of pressure derivatives for the different types of vibrational modes observed indicates that each molecular unit readjusts internally to pressure induced volume changes and the some of the chemical bonds between the units are significantly affected.     

Theoretical calculation of equilibrium Mg isotope fractionation between silicate melt and its vapor

Isotope fractionation during the evaporation of silicate melt and condensation of vapor has been widely used to explain various isotope signals observed in lunar soils, cosmic spherules, calcium–aluminum-rich inclusions, and bulk compositions of planetary materials. During evaporation and condensation, the equilibrium isotope fractionation factor (α) between high-temperature silicate melt and vapor is a fundamental parameter that can constrain the melt’s isotopic compositions. However, equilibrium α is difficult to calibrate experimentally. Here we used Mg as an example and calculated equilibrium Mg isotope fractionation in MgSiO₃ and Mg₂SiO₄ melt–vapor systems based on first-principles molecular dynamics and the high-temperature approximation of the Bigeleisen–Mayer equation. We found that, at 2500 K, δ²⁵Mg values in the MgSiO₃ and Mg₂SiO₄ melts were 0.141?±?0.004 and 0.143?±?0.003‰ more positive than in their respective vapors. The corresponding δ²⁶Mg values were 0.270?±?0.008 and 0.274?±?0.006‰ more positive than in vapors, respectively. The general \(\alpha - T\) equations describing the equilibrium Mg α in MgSiO₃ and Mg₂SiO₄ melt–vapor systems were: \(\alpha_{{{\text{Mg}}\left( {\text{l}} \right) - {\text{Mg}}\left( {\text{g}} \right)}} = 1 + \frac{{5.264 \times 10^{5} }}{{T^{2} }}\left( {\frac{1}{m} - \frac{1}{{m^{\prime}}}} \right)\) and \(\alpha_{{{\text{Mg}}\left( {\text{l}} \right) - {\text{Mg}}\left( {\text{g}} \right)}} = 1 + \frac{{5.340 \times 10^{5} }}{{T^{2} }}\left( {\frac{1}{m} - \frac{1}{{m^{\prime}}}} \right)\), respectively, where m is the mass of light isotope ²⁴Mg and m′ is the mass of the heavier isotope, ²⁵Mg or ²⁶Mg. These results offer a necessary parameter for mechanistic understanding of Mg isotope fractionation during evaporation and condensation that commonly occurs during the early stages of planetary formation and evolution.     

Some geobarometers involving cordierite in the FeO-Al2O3-SiO2(±H2O) system: refinements,thermodynamic calibration,and applicability in granulite facies rocks

Five geobarometers involving cordierite have been formulated for quantitative pressure sensing in high grade metapelites. The relevant reactions in the FeO-Al₂O₃-SiO₂ (±H₂O) system are based on the assemblages (A) cordierite-garnet-sillimanite-quartz, (B) cordierite-spinel-quartz, (C) cordierite-garnet-spinel-sillimanite, (D) cordierite-garnet-orthopyroxene-quartz and (E) cordierite-orthopyroxene-sillimanite-quartz. Application of the barometric formulations to a large number of granulite grade rocks indicates that the cordierite-garnet-sillimanite-quartz equilibrium is widely applicable and registers pressures which are in good agreement with the “consensus” pressure estimates. The dispersion in the computed P values, expressed as one standard deviation, is within ±1.2 kbar. The geobarometers (B) and (C) also yield pressures which are reasonable and compare well with those computed from equilibrium (A). The estimated pressures from (D) and (E), both involving orthopyroxene, are at variance with these estimates. It has been argued that the discrepancy in pressures obtained from these geobarometers stems from an inadequate knowledge of activity-composition relations and/or errors in input thermodynamic data of aluminous orthopyroxene. The convergence of pressure values estimated from the barometric formulations, especially (A), (B) and (C), implies that the present formulations are more dependable than the existing formulations and are also capable of setting limits on P values in response to varying $$\begin{gathered} {\text{1/2Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} \hfill \\ {\text{ = 1/3Fe}}_{\text{3}} {\text{Al}}_{\text{2}} {\text{Si}}_{\text{3}} {\text{O}}_{{\text{12}}} {\text{ + 2/3Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{ + 5/6SiO}}_{\text{2}} {\text{. (A)}} \hfill \\ {\text{1/2Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} {\text{ = FeAl}}_{\text{2}} {\text{O}}_{\text{4}} {\text{ + 5/2SiO}}_{\text{2}} {\text{. (B)}} \hfill \\ {\text{Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} {\text{ + FeAl}}_{\text{2}} {\text{O}}_{\text{4}} \hfill \\ = {\text{Fe}}_{\text{3}} {\text{Al}}_{\text{2}} {\text{Si}}_{\text{3}} {\text{O}}_{{\text{12}}} {\text{ + 2Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{. (C)}} \hfill \\ {\text{1/2Fe}}_{\text{2}} {\text{Al}}_{\text{4}} {\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} {\text{ + Fe}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} \hfill \\ = {\text{Fe}}_{\text{3}} {\text{Al}}_{\text{2}} {\text{Si}}_{\text{3}} {\text{O}}_{{\text{12}}} {\text{ + 3/2SiO}}_{\text{2}} .{\text{ (D)}} \hfill \\ {\text{1/2Fe}}_{\text{2}} {\text{Al}}{}_{\text{4}}{\text{Si}}_{\text{5}} {\text{O}}_{{\text{18}}} \hfill \\ = 1/2{\text{Fe}}_{\text{2}} {\text{Si}}_{\text{2}} {\text{O}}_{\text{6}} {\text{ + Al}}_{\text{2}} {\text{SiO}}_{\text{5}} {\text{ + 1/2SiO}}_{\text{2}} .{\text{ (E)}} \hfill \\ \end{gathered}$$ . The present communication addresses the calibration, applicability and reliability of these barometers with reference to granulite facies metapelites.     

Optical absorption and mössbauer spectra of purple and green yoderite,a kyanite-related mineral

Mössbauer and polarized optical absorption spectra of the kyanite-related mineral yoderite were recorded. Mössbauer spectra of the purple (PY) and green yoderite (GY) from Mautia Hill, Tanzania, show that the bulk of the iron is Fe³⁺ in both varieties, with Fe²⁺/(Fe²⁺+Fe³⁺) ratios near 0.05. Combining this result with new microprobe data for PY and with literature data for GY gives the crystallochemical formulae: $$\begin{gathered} ({\text{Mg}}_{{\text{1}}{\text{.95}}} {\text{Fe}}_{{\text{0}}{\text{.02}}}^{{\text{2 + }}} {\text{Mn}}_{{\text{0}}{\text{.01}}}^{{\text{2 + }}} {\text{Fe}}_{{\text{0}}{\text{.34}}}^{{\text{3 + }}} {\text{Mn}}_{{\text{0}}{\text{.07}}}^{{\text{3 + }}} {\text{Ti}}_{{\text{0}}{\text{.01}}} {\text{Al}}_{{\text{3}}{\text{.57}}} )_{5.97}^{[5,6]} \hfill \\ {\text{Al}}_{{\text{2}}{\text{.00}}}^{{\text{[5]}}} [({\text{Si}}_{{\text{3}}{\text{.98}}} {\text{P}}_{{\text{0}}{\text{.03}}} ){\text{O}}_{{\text{18}}{\text{.02}}} ({\text{OH)}}_{{\text{1}}{\text{.98}}} ] \hfill \\ \end{gathered}$$ and PY and $$\begin{gathered} ({\text{Mg}}_{{\text{1}}{\text{.98}}} {\text{Fe}}_{{\text{0}}{\text{.02}}}^{{\text{2 + }}} {\text{Mn}}_{{\text{< 0}}{\text{.001}}}^{{\text{2 + }}} {\text{Fe}}_{{\text{0}}{\text{.45}}}^{{\text{3 + }}} {\text{Ti}}_{{\text{0}}{\text{.01}}} {\text{Al}}_{{\text{3}}{\text{.56}}} )_{6.02}^{[5,6]} \hfill \\ {\text{Al}}_{{\text{2}}{\text{.00}}}^{{\text{[5]}}} [({\text{Si}}_{{\text{3}}{\text{.91}}} {\text{O}}_{{\text{17}}{\text{.73}}} {\text{(OH)}}_{{\text{2}}{\text{.27}}} ] \hfill \\ \end{gathered}$$ for GY. The Mössbauer spectra at room temperature contain one main doublet with isomer shifts and quadrupole splittings of 0.36 (PY), 0.38 (GY) and 1.00 (PY), 0.92 (GY) mm s^?1, respectively. These values correspond to Fe³⁺ in six or five-fold coordination. The doublet components have anomalously large half widths indicating either accomodation of Fe³⁺ in more than one position (e.g., octahedraA1 and five coordinatedA2) or the yet unresolved superstructure. Besides strong absorption in the ultraviolet (UV) starting from about 25,000 cm^?1, the polarized optical absorption spectra are dominated by strong bands around 16,500 and 21,000 cm^?1 (PY) and a medium strong band at around 13,800 cm^?1 (GY). Position and polarization of these bands, in combination with the UV absorption, explain the colour and pleochroism of the two varieties. The bands in question are assigned to homonuclear metal-to-metal charge transfer transitions: Mn²⁺(A1) Mn³⁺(A1′) ? Mn³⁺(A1) Mn²⁺(A1′) and Mn²⁺(A1) Mn³⁺(A2 ? Mn³⁺(A1) Mn²⁺(A2) in PY and Fe²⁺(A1) Fe³⁺(A1′) ? Fe³⁺(A1) Fe²⁺(A1′) in GY. The evidence for homonuclear Mn²⁺ Mn³⁺ charge transfer (CTF) is not quite clear and needs further study. Heteronuclear FeTi CTF does not contribute to the spectra. In PY, additional weak bands were resolved at energies around 17,700, 18,700, 21,000, and 21,900 cm^?1 and assigned to Mn³⁺ in two positions. Weak bands around 10,000 cm^?1 in both varieties are assigned to Fe²⁺ spin-alloweddd-transitions. Very weak and sharp bands, around 15,400, 16,400, 21,300, 22,100, 23,800, and 25,000 cm^?1 are identified in GY and assigned to Fe³⁺ spin-forbiddendd-transitions.     

Kanonaite, (Mn 0.76 3+ Al0.23Fe 0.02 3+ )[6]Al[5][O¦SiO4], a new mineral isotypic with andalusite

Kanonaite forms rare porphyroblasts up to 12mm long in a gahnite— Mg-chlorite — coronadite — quartz schist occurring near Kanona, Zambia. The composition is (microprobe analysis): SiO₂ 32.2, Al₂O₃ 33.9, Mn as Mn₂O₃ 32.2, Fe₂O₃ 0.66, ZnO 0.13, MgO 0.04, BaO 0.04, TiO₂ 0.01, CaO 0.01, PbO 0.01, CuO 0.01, total 99.21, corresponding to $$\left( {{\text{Mn}}_{{\text{0}}{\text{.76}}}^{{\text{3 + }}} {\text{Al}}_{{\text{0}}{\text{.23}}} {\text{Fe}}_{{\text{0}}{\text{.015}}}^{{\text{3 + }}} } \right)_{1.005}^{\left[ 6 \right]} {\text{AL}}_{1.00}^{\left[ 5 \right]} \left[ {{\text{O}}_{{\text{1}}{\text{.00}}} |{\text{Si}}_{{\text{0}}{\text{.99}}} {\text{O}}_{{\text{4}}{\text{.00}}} } \right]$$ The mineral is greenish black, strongly pleochroic with X(∥a) yellow green, Y(∥b) bluish green, Z(∥c) deep golden yellow, biaxial positive, with 2V = 53°(3°), α = 1.702, β = 1.730, γ = 1.823. Vickers microhardness (100 gram load) ranges between 906 and 1017kp/mm². The structure is orthorhombic, isotypic with andalusite, space group Pnnm, a = 0.7953(2), b = 0.8038(2), c = 0.5619(2) nm, V = 0.3592(1) nm³, a/b = 0.9895(3), c/b = 0.6990(3), S.G.(_x) = 3.395 g/cm³, Z = 4. The strongest X-ray powder lines are (d in nm, I, hkl):0.5669, 100, 110; 0.4590, 75, 011 and 101; 0.3577, 90, 120 and 210; 0.2827, 94, 220; 0.2517, 90, 310 and 112; 0.2212, 83, 320, 122 and 212. Comparison of the intensities of 373 observed X-ray reflections with those calculated for several models of Mn³⁺-distribution indicates octahedral coordination of all or most of the manganese present. Interpretation of magnetic measurements (μ_eff = 3.15B.M. per Mn atom at 25 ° C) indirectly supports octahedral coordination of Mn³⁺. The name of the mineral is for Kanona, a town near the type locality. The name is proposed for the end member Mn^{3+ [6]}Al^[5][O¦SiO₄] and for members of the solid-solution series towards andalusite with octahedral Mn³⁺>Al. The presently described mineral may be referred to as aluminian kanonaite.