Volume 80, Issue 1-2

An in-situ study of cubanite, CuFe 2 S 3 , was performed in a diamond-anvil cell using Mossbauer spectroscopy and energy-dispersive X-ray diffraction at room temperature and pressures up to 5 GPa. Mossbauer spectra of orthorhombic cubanite show a single Fe site with rapid electron transfer between Fe 2+ and Fe 3+ , a hyperfine magnetic field that is relatively insensitive to pressure, and a center shift that decreases with pressure because of increasing covalency. A phase transition occurs above 3.3 GPa that involves a change from the orthorhombic cubanite structure to a derivative of the hexagonal NiAs (B8) structure, with a zero-pressure volume decrease of 29%. The large difference in volume is caused by a change from tetrahedral to octahedral coordination and a significant shortening of metal-metal bonds. Volume-compression data were fitted to a second-order Birch- Murnaghan equation of state (K′ 0 = 4) with the results K 0 = 64 ± 3 GPa (orthorhombic phase) and K 0 = 157 ± 16 GPa (high-pressure phase). Mossbauer data of the high-pressure phase indicate a single Fe site with no magnetic ordering and a valence intermediate between Fe 2+ nd Fe 3+ . Consideration of likely ordering patterns in the high-pressure phase indicates that localized electron transfer could occur along face-shared pairs of Fe octahedra, and extended electron delocalization could occur along paths formed by faceand edge-shared octahedra.

From in-situ structure and site occupancy refinements of a synthetic orthopyroxene, (Mg 0.75 Fe 0.25 ) 2 Si 2 O 6 , by single-crystal X-ray diffraction at 296,1000, 1100,1200, and 1300 K (with reversals), two significant phenomena were observed: (I) with increasing temper­ature, the silicate B chain straightens much faster, especially above 1200 K, and becomes straighter than the A chain at 1300 K, with the 03-03-03 angles of the A and B chains being 170.8 and 173.1°, respectively; and (2) the In K D values (where K D = Fe M1 Mg M2 / Fe M2 Mg M1 ) vary linearly with MT (K) between 1000 and 1200 K, but nonlinearly between 1200 and 1300 K. The drastic straightening of the silicate chains at 1300 K is accompanied by the attainment of very similar configurations of the SiO 4 tetrahedra, SiA and SiB, in shapes, sizes, and out-of-plane tiltings. The rigid-body thermal vibration analysis suggests that the librational motion of the SiA tetrahedron (8. I° 2 ) is slightly larger than that of SiB (5.2° 2 ) at 296 K but becomes considerably smaller (48.0° 2 ) than that of SiB (77.3° 2 ) at 1300 K. The Ml octahedron remains nearly regular from 296 to 1300 K, whereas the M2 coordination changes from sixfold (296 K) to sevenfold (1200 K) and to sixfold (1300 K) because of the bridging 03B moving out and 03B′ moving into the coordination sphere (r ≤ 3.0 Å) at elevated temperatures. At 1300 K, the highly distorted M2 octahedron shares two edges with SiA and SiB, rather than one with SiA at 296 K, and the structure has all features of the high-temperature orthopyroxene phase predicted by Pannhorst (1979), which, in fact, is a transitional structural state between orthopyroxene and the so-called protoenstatite. On the basis of the configurations of the silicate chains in Mg- and Fe-rich orthopyroxenes at high temperature, an explanation is given as to why Mg-rich orthopy­roxenes tend to transform to the protoenstatite structure with increasing temperature, whereas Fe-rich ones tend to transform to a clinopyroxene with space group C2/c. The anomalous behavior of the Fe-Mg order-disorder above 1200 K is attributed to the exis­tence of the transitional state, particularly the changing charge distribution around the M2 site.

Monazite and xenotime, the RE(PO 4 ) dimorphs, are the most ubiquitous rare earth (RE) minerals, yet accurate structure studies of the natural phases have not been reported. Here we report the results of high-precision structure studies of both the natural phases and the synthetic RE(PO 4 ) phases for all individual stable rare earth elements. Monazite is monoclinic, P2 1 /n, and xenotime is isostructural with zircon (space group I4 1 /amd). Both atomic arrangements are based on [001] chains of intervening phosphate tetrahedra and RE polyhedra, with a REO 8 polyhedron in xenotime that accommodates the heavy lanthanides (Tb-Lu in the synthetic phases) and a REO 9 , polyhedron in monazite that preferentially incorporates the larger light rare earth elements (La-Gd). As the struc­ture “transforms” from xenotime to monazite, the crystallographic properties are com­parable along the [001] chains, with structural adjustments to the different sizes of RE atoms occurring principally in (001). There are distinct similarities between the structures that are evident when their atomic arrangements are projected down [001]. In that projection, the chains exist in (100) planes, with two planes per unit cell. In monazite the planes are offset by 2.2 Å along [010], relative to those in xenotime, in order to accommodate the larger light RE atoms. The shift of the planes in monazite allows the RE atom in that phase to bond to an additional O2′ atom to complete the REO 9 polyhedron.

A homogeneous megacryst of schorlomite was investigated to determine the valence states of Fe and Ti and the crystallographic sites occupied by these elements. The chemical composition of the specimen was analyzed by electron microprobe, wet-chemical analysis, FTlR, and INAA. The results from X-ray absorption near-edge structure spectroscopy (XANES) are consistent with exclusively Ti 4+ occupying the octahedral site only. The tetrahedral site is deficient in Si and the results of low-temperature 57 Fe Mossbauer spec­troscopy indicate that the remainder of the site is occupied by Fe 3+ and substantial Fe 2+ . A spin-allowed intensified crystal-field transition of [4] Fe 2+ is present in the near-infrared spectrum. The optical absorption spectrum is dominated by an intense band centered at 500 nm with a full width of 8000 cm -1 at half maximum peak height; this band is assigned to an Fe 2+ -Ti 4+ intervalence charge transfer transition between ,[4] Fe 2+ and [6] Ti. The cation site occupancies in this specimen of schorlomite can be expressed by the following formula: {Ca 2.866 Mg 0.800 Na 0.038 Mn 0.019 }Σ 3.003 [Ti 4+ 1.058 Fe 3+ 0.631 Al 0.137 Fe 2+ 0.057 Mg 0.055 Zr 0.039 V 3+ 0.014 Mn 0.013 ] Σ 2 . 004 - (Si 2.348 Fe 3+ 0.339 Fe 2+ 0.311 [4H] 0.005 ) Σ3.003 O 12 .

The minerals zunyite, Al 13 Si 5 O 20 (OH 9 F) 18 Q 5 and harkerite, Ca 24 Mg g [AlSi 4 (O 5 OH) 16 ] 2 - (CO 3 ) 8 (BO 3 ) 8 (H 2 O 9 Q) 5 have been studied by means of solid-state 27 Al MAS NMR. Zunyite contains Si 5 O 16 pentamers and harkerite contains AlSi 4 (O 5 OH) 16 pentamers. These pen- tameric groups are unique because their T-O-Si angles are almost 180°. Chemical analysis of the zunyite sample shows that it has excess Al: the Al-Si ratio is 2.9, compared with the ideal of 2.6. High-speed spinning 27 Al MAS NMR spectra (11-13 kHz) showed two more signals than the spectra obtained by Kunwar et al. (1984). The signal with δ iso = 46.8 ± 0.5 ppm represents the excess Al, which enters the central Si1 site of the Si 5 O 16 pentamer. This assumption is confirmed by the fact that δ iso of Al in the AlSi 4 (O,OH) 16 pentamer in harkerite is 44 ± 1 ppm. Additional proof comes from comparing the electrostatic energy and the quadrupole interaction of Al in either a Si1 or Si2 configuration. The Al site in the pentamers of zunyite and harkerite can be considered as a q 4 (4Si) site. In this case, Al-O-Si angles are correlated with the 27 Al chemical shift (Lippmaa et al., 1986). This correlation holds well for the 27 Al data for harkerite. The value for zunyite indicates that the structure adapts to the incorporation of Al in the Sil site by a narrowing of the Al-O-Si angle to 171 ± 2.5°. The lower limit of the chemical shift range for Al in framework aluminosilicates is decreased by 12 ppm, from 55.8 ppm for mordenite to 44 ppm for harkerite.

Impedance spectroscopy helps distinguish the contributions that grain interiors and grain boundaries make to electrical resistance of silicate minerals and rocks. The technique also distinguishes the low-frequency response due to the presence of instrument electrodes. We measured olivine, orthopyroxene, clinopyroxenes, and both natural and synthetic clinopyroxenite. Measurements were made at 1 bar, from 750 to 1150 °C, and over a fre­quency range from < 10 -4 to > 10 6 Hz; some measurements were also made at 300-850 °C and 10-20 kbar. The grain-interior response lies at highest frequency, the sample- electrode response at low frequencies, and the grain boundary response at mid-frequencies. Grain interiors show as semicircular impedance arcs when plotted on the complex plane, and sample-electrode responses of hot single crystals and of hot dry rocks are exhibited as depressed arcs. In comparison, monofrequency measurements contain no information to identify the source of the response; at 1 kHz they detect only the resistance sum of grain interiors and grain boundaries and at low frequency (≤ 1 Hz) are likely to sense all three components. The major experimental problem is to find electrodes that make good contact with the sample and that are stable with time. The effect of pressure (10 kbar, 300-800 °C) is to diminish the resistance associated with grain boundaries and the sample-electrode interface, in the laboratory and presumably in nature. Monofrequency measurements at 1 bar may underestimate the conductivity of rocks at similar temperature but higher pressure. A network of electrical elements is presented for use in interpreting impedance spectra and conductive paths in hot or cold, wet or dry, minerals and rocks at any pressure. In dry rocks, a series network path predominates; in wet rocks, aqueous pore fluid and crystals both conduct. Finite resistance across the sample-electrode interface is evidence that elec­tronic charge carriers are present at the surface, and presumably within, the silicate min­erals and rocks measured.

Dozyite is a new mineral species involving regular interstratification of trioctahedral serpentine and trioctahedral chlorite units in a 1:1 ratio. It occurs as colorless crystals in an altered skarn adjacent to the Ertsberg East Cu, Au, and Ag mine in central Irian Jaya, Indonesia. The name is after Jean Jacques Dozy, the Dutch geologist who discovered and named the Ertsberg ore province in 1936. Unit-cell parameters are a = 5.323(3), b = 9.214(9), c = 21.45(2) A, β= 94.43(6)°, and V = 1049(2) Å 3 .It has space group Cm. There is a 21-Å periodicity in the 00l and in most other reflections where k = 3n. The reflections where k ≠3n are continuously streaked. Excellent regularity of alternation of the component serpentine and chlorite units is indicated by the coefficient of variation CV = 0.26 for the 001 reflections. A simplified ideal bulk composition is (Mg 7 Al 2 )(Si 4 Al 2 )O 15 - (OH) 12 with Z = 2, halfway between the compositions of the closely associated discrete chlorite (clinochlore) and the discrete serpentine (amesite). We believe the components in this occurrence of dozyite are clinochlore and amesite and that the interstratification was formed during the conversion of early clinochlore to amesite by Al metasomatism. The structure of dozyite contains a la chlorite unit followed by a serpentine 1:1 layer that is in the same position that the lower tetrahedral sheet of a repeating chlorite unit would occupy in the one-layer monoclinic Iaa-2 chlorite polytype, but rotated 180° so that the octahedral cations alternate I,I,II per 21 Å. The next chlorite unit follows with zero shift of its lower sixfold rings relative to those of the serpentine 1:1 layer. A second occurrence of dozyite has been recognized in a Cr-rich serpentinite from the Wood Chrome mine in Lancaster County, Pennsylvania. It represents a different polytype, with β= 90°, and a different composition relative to the Ertsberg material.

Pelitic schists in the contact aureole of the Iate-Alpine Vedrette di Ries granodiorite show an epitaxial replacement of kyanite by Fe-rich staurolite. The reaction is observed within nodules of fine-grained staurolite crystals (100-200 μm) containing oriented la­mellae of kyanite. Microprobe, backscattered electron microscopy (BSEM), and high-resolution transmis­sion electron microscopy (HRTEM) investigations demonstrate the presence of submicro- scopic intergrowths of staurolite and kyanite. The orientation of both phases, a st ‖b Ky and cst‖c Ky ,is in agreement with previous observations. Individual staurolite and kyanite la­mellae are between one unit cell and some tens of unit cells wide. Detailed observation of the termination of staurolite lamellae within kyanite suggests that the replacement occurs by a solid-state reaction, in which the primary cubic close-packed O framework of kyanite is preserved. The epitaxial replacement implies a reorganization of the site occupancy accompanied by addition of Fe + Mg + H and removal of Si + Al from the reaction site. The epitaxial replacement of metastable kyanite occurred during contact metamorphism within the andalusite stability field, under temperature and pressure conditions of 500-550 °C and >2.9 kbar, respectively.

Optically anisotropic and isotropic haiiyne has been observed in phonolite from Niang- niang Hill, Jiangsu Province, southeastern China. Selected-area electron diffraction (SAED) and transmission electron microscopy (TEM) studies show the anisotropic haiiyne has a one-dimensional superstructure with a period of 6d I10sub along one of the <110) directions of the ideal cubic haiiyne subcell. The superstructure may be considered as a periodic arrangement of (110) layer domains with P23 symmetry along [110]. Neighboring domains are related by a twin relationship. The unit-cell parameters for this haiiyne superstructure are thus metrically orthorhombic, with a sup = 2d 1̄ 10sub ,b sup = 6d 110sub , C sup = c sub , but the most likely space group is monoclinic, Pn. SAED patterns and TEM images of two opti­cally isotropic haiiyne samples show they have either a three-dimensional superstructure or no superstructure. The average structure of the isotropic haiiyne with the superstructure is cubic, even though it is composed of twin domains with one-, two-, and three-dimen­sional superstructures. The superstructure may have formed by ordering Of[Na 3 Ca(SO 4 )] 3+ and [(Na,K) 4 (OH)] 3+ clusters in the cubo-octahedral cages of the haüyne framework during an arrested phase transition from P4̄3n to P23 symmetry.

H 2 O solubility has been determined in haplogranitic melts (system SiO 2 -NaAlSi 3 O 8 - KAlSi 3 O 8 , Qz-Ab-Or) in the range 0.5-8 kbar and 800-1350 °C Three types of starting materials were used: dry glass cylinders, prehydrated glass pieces, or dry glass blocks surrounded by glass powder. All the starting materials gave consistent results. The H 2 O contents of the glasses were determined by Karl-Fischer titration. Dissolved H 2 O was demonstrated to be distributed homogeneously throughout the isobarically quenched melts (glasses) using infrared spectroscopy. The compositional dependence of H 2 O solubility was mainly determined at 0.5 kbar, 900 and 1000 °C; I kbar, 850 0 C; and 4.8 kbar, 800 0 C. Seventeen compositions containing 25, 35, or 45 wt% normative Qz and with various Or/(Or + Ab) ratios (0.86-0.09, Ab and Or expressed as normative weight percent) were investigated. At 0.5 kbar, H 2 O sol­ubility was little affected by the anhydrous composition. By contrast, molar H 2 O solubility in aluminosilicate melts was significantly dependent upon anhydrous composition between I and 5 kbar. The highest solubility values were obtained for the most Ab-rich melts. This alkali effect has important implications for the physical and chemical properties of granitic melts. The effect of pressure (P) on H 2 O solubility at P ≥ 3 kbar is greater than that reported in previous studies. Between 3 and 8 kbar at 800 °C, there is a (nearly linear) positive correlation between P and H 2 O solubility. The effect of temperature (T) on H 2 O solubility was investigated for a composition Qz 28 Ab 38 Or 34 (normative weight percent) in the P-T range 0.5-8 kbar and 800-1350 °C. Water solubility ranged from retrograde (with increas­ing T) at P ≤ 4 kbar through temperature independence at approximately 4.5 kbar to prograde at P = 5 kbar. Calculated H 2 O solubilities using the model of Burnham and Nekvasil (1986) are slightly high at 0.5 kbar and significantly low at 5 kbar, compared with the experimental data. This implies that calculated H 2 O activities for haplogranitic systems using the H 2 O content of the melt may be overestimated at high pressure (P ≥ 5 kbar). Using the ther­modynamic model of Silver and Stolper (1985) and assuming a proportion of molecular H 2 O and OH groups close to that defined for albite melts by Silver and Stolper (1989), we found that the partial molar volume of H 2 O in a melt with a composition Qz 28 Ab 38 Or 34 has to be close to 10-12 cm 3 /mol to obtain a good agreement between the calculated and the experimentally determined H 2 O solubility curves in the pressure range 1-8 kbar at 900 °C.

The effects of packing materials on H infiltration into sample capsules during pistoncylinder experiments has been examined using samples of oxalic acid dihydrate, silver oxalate, and pure carbon dioxide. Mass scans revealed that as much as several mole percent carbon monoxide, accompanied by H 2 O, can result from a reaction between H and carbon dioxide in experiments using all three reactants. A measurable 0 isotopic shift was detected in carbon dioxide samples where carbon monoxide was detected. Traditional means for measuring C-O-H volatile composition at the end of piston-cylinder experiments will not identify this chemical change in the volatile composition and may result m erroneous determmation of fluid composition. Glass and witherite packing materials effectivelymaintained a low f H2 around the samples.

A thermochcmical model of the activities of species in carbonate-rich melts would be useful in quantifying chemical equilibria between carbonatite magmas and vapors and in extrapolating liquidus equilibria to unexplored PTX. A regular-solution model of Ca-rich carbonate melts is developed here, using the fact that they are ionic liquids, and can be treated (to a first approximation) as interpenetrating regular solutions of cations and of anions. Thermochemical data on systems of alkali metal cations with carbonate and other anions are drawn from the literature; data on systems with alkaline earth (and other) cations and carbonate (and other) anions are derived here from liquidus phase equilibria. The model is validated in that all available data (at 1 kbar) are consistent with single values for the melting temperature and heat of fusion for calcite, and all liquidi are con­sistent with the liquids acting as regular solutions. At 1 kbar, the metastable congruent melting temperature of calcite (CaCO 3 ) is inferred to be 1596 K, with ΔH̅ fus (Calcite) = 31.5 ± 1 kJ/mol. Regular solution interaction param­eters (IV) for Ca 2+ and alkali metal cations are in the range -3 to -12 kj/mol 2 ; IV for Ca 2+ -Ba 2+ is approximately -11 kJ/mol 2 ; W for Ca 2+ -Mg 2+ is approximately -40 kJ/ mol 2 , and W for Ca 2+ -La 2+ is approximately +85 kJ/mol 2 . Solutions of carbonate and most anions (including OH - , F - , and SO 2- 4 ) are nearly ideal, with W between O (ideal) and -2.5 kJ/mol 2 . The interaction of carbonate and phosphate ions is strongly nonideal, which is consistent with the suggestion of carbonate-phosphate liquid immiscibility. Interaction of carbonate and sulfide ions is also nonideal and suggestive of carbonate-sulfide liquid immiscibility. Solution of H 2 O, for all but the most H 2 O-rich compositions, can be mod­eled as a disproportionation to hydronium (H 3 O + ) and hydroxyl (OH - ) ions with W for Ca 2+ -H 3 O + ≈ 33 kJ/mol 2 . The regular-solution model of carbonate melts can be applied to problems of carbonatite magma + vapor equilibria and of extrapolating liquidus equilibria to unstudied systems. Calculations on one carbonatite (the Husereau dike, Oka complex, Quebec, Canada) show that the anion solution of its magma contained an OH - mole fraction of ∼0.07, although the vapor in equilibrium with the magma had P(H 2 O) = 8.5 × P(CO 2 ). F in carbonatite systems is calculated to be strongly partitioned into the magma (as F - ) relative to coexisting vapor. In the Husereau carbonatite magma, the anion solution contained an F mole fraction of ∼6 × 10 -5 . Calcite and anhydrite may be present on the surface of Venus, but they would not be molten at ambient surface temperature (660-760 K) because the minimum melt temper­ature (eutectic) for the calcite + anhydrite system is calculated to be 1250 K. The Venus atmosphere contains 5 ppb HF, which implies that the anion solution of a carbonate-rich magma in equilibrium with the atmosphere would contain a F - mole fraction of ∼7 × 10 -3 or about 0.1 wt%. Although this proportion of F is much enriched compared with the atmosphere, it would have little effect on phase relations of the carbonatite.

Temperatures based on the composition of calcite coexisting with dolomite (calcite + dolomite thermometry) range from 475 to 600 °C for 63 marbles from the Llano uplift of central Texas. The highest temperatures, ∼600 °C, were obtained by carefully reintegrating calcite containing exsolved lamellae of dolomite. In some cases, these high temperatures were determined for marbles that contain an isobarically invariant assemblage consisting of calcite + dolomite + tremolite + diopside + forsterite. At a pressure of 3 kbar, these five minerals are stable at 630 °C and X CO2 = 0.62. In contrast, relatively low calcite + dolomite temperatures of 475-480 °C were obtained for marbles containing the assemblage calcite + dolomite + tremolite + talc. This talc-bearing assemblage is stable at ≤475 °C, depending on fluid composition, at a pressure of 3 kbar. Additional isobarically univariant equilibria are stable at intermediate temperatures (generally between 535 and 630 °C), and these are also generally consistent with results obtained from calcite + dolomite thermometry. Agreement between the calcite + dolomite temperatures and those inferred from silicate + carbonate equilibria in the marbles indicates that the temperatures generally reflect peak conditions of metamorphism, although some resetting has occurred. The mar­bles having the highest calcite-dolomite temperatures, as well as those containing the high- temperature isobarically invariant assemblage, are generally found close to post-tectonic pluton contacts, indicating that some of the amphibolite-facies assemblages are related to the emplacement of the granitic intrusions. Relatively low temperatures are recorded with­in approximately 2 km of pluton contacts, suggesting a possible thermal aureole. However, relatively high temperatures of 550-600 °C are recorded in marbles that are not spatially related to post-tectonic granites (>2 km from pluton contacts). These temperatures may be relicts of an earlier metamorphic event, although they could be related to granites that were emplaced above or below the current erosional level.

Newly discovered orthopyroxene and clinopyroxene megacrysts (OPMs and CPMs) in the Labrieville massif provide additional constraints on megacryst origins and their sig­nificance in massif anorthosite petrogenesis. Pyroxene megacrysts (∼ 1-15 cm across) oc­cur throughout the anorthosite, some in layered zones (as cumulates?) and others as scat­tered crystals that form subophitic intergrowths with plagioclase; some also contain tabular plagioclase inclusions. In many cases, pyroxene megacrysts are intergrown with ilmenite and locally occur within meter-sized masses of ilmenite. Plagioclase and ilmenite are the dominant exsolution products in OPMs, whereas CPMs contain additional lamellae of orthopyroxene. Biotite consistently occurs with megacrysts, locally as an integral portion of exsolution lamellae. Microprobe spot compositions of the megacrystic pyroxenes (OPM ∼ En 58-66 , CPM ∼ Wo 45 En 40 Fs 15 ) are essentially the same as those of smaller matrix pyroxenes in the same outcrops. In contrast, three bulk OPMs (4.4-5.9 wt% Al 2 O 3 ) and one CPM (6.4 wt% Al 2 O 3 ) are enriched in Ca, Na, Al, and Ti, compatible with observed exsolution products. Plagio­clase lamellae (An 34-53 ) are more calcic than matrix plagioclase (An 32-38 ) but are the most sodic recognized in any massif anorthosite thus far. Field characteristics of megacrysts indicate that they formed contemporaneously with the other minerals in the same outcrops. Furthermore, bulk OPMs have X Mt similar to matrix pyroxene (or are slightly more Fe-rich). In comparison, OPMs from the Saint Urbain massif are more magnesian (∼En 66-72 ), whereas lamellae and matrix plagioclase are more anorthitic (∼An 55-75 and ∼An 40 , respectively). These systematic differences be­tween the two massifs suggest that the En content of OPMs and the An content of plagio­clase (matrix and lamellae) are strongly controlled by local magma composition. Collec­tively, these results can be explained best by the in-situ growth of megacrysts, together with or after adjacent plagioclase. The megacrysts bulk compositions can be adequately accounted for by a rapid growth mechanism. The common association of OPMs with ilmenite (and biotite) is noteworthy and suggests that oxidizing conditions played an im­portant role in controlling the properties and compositions of the megacrysts. The impli­cation is that these megacrysts are not relics of prior high-pressure pyroxene fractionation and provide no demonstrable link to a basaltic parent magma for this massif.

Strongly zoned amphibole formed oftitanian fluor-richterite cores with 2.24% TiO 2 and margins oftitanian ferro-richterite with 6.65% TiO 2 is described from a basaltic andesite. The cores are richer in F, AI, Mg, and Ca than the edges, which are richer in OH, Ti, Fe, Na, and K. Similar sodic-calcic amphiboles are widely known to appear in lamproites, but typically they contain 5% or more K 2 O, whereas the present sample reaches a maximum of 1.18% K 2 0 at the margin. As is characteristic of titanian potassium richterite, there is Ti in tetrahedral coordination and a pink to yellow pleochroism. Ba, Sr, Li, B, CI, Be, Nb, La, Ce, Pr, Y, V, and Zr contents are given. Zoned augite with Ca 42 Mg42Fe 16 cores and Ca 43 Mg 36 Fe 22 rims is also present, plus chloritic material and feldspar.

Ungarettiite is a new amphibole species from the Hoskins mine, near Grenfell, New South Wales, Australia. It occurs with Mn-bearing oxides, silicates, and carbonates in a stratiform schist associated with metajasper, metabasalt, and metasiltstone. Ungarettiite is brittle, H = 6, D meas = 3.52 g/cm 3 , D calc = 3.45 g/cm 3 . In plane-polarized light, it is strongly pleochroic, X = orange red, Y ~ Z = very dark red; X Λ a = -2° (in β acute), Y = b, Z Λ C= 17° (in β obtuse), with absorption X < Y ≤ Z. Ungarettiite is biaxial positive, α = 1.717(2), β = 1.780(4), γ= 1.800(2); 2V= 51(2)°, dispersion r < ν. Ungarettiite is monoclinic, space group C2/m, a = 9.89(2), b = 18.04(3), c = 5.29(1) Å, β = 104.6(2)°, V = 912(1) Å 3 , Z = 2. The strongest ten X-ray diffraction lines in the powder pattern are [d(I, hkl )]: 2.176(10,171), 3.146(9,310), 2.544(9,2̅02), 1.447(9,3̅.11.1), 3.400(8,131), 1.656(8,461), 8.522(7,110), 2.299(7,1̅71), 2.575(6,241), and 2.047(6,202). Analysis by a combination of electron microprobe and crystal-structure refinement gives SiO 2 50.66, Al 2 O 3 0.04, TiO 2 0.03, Fe 2 O 3 0.50, FeO 0.00, Mn 2 O 3 24.35, MnO 12.42, MgO 1.46, ZnO 0.10, CaO 0.18, Na 2 O 9.13, K 2 O 0.76, F not detected, sum 99.63 wt%. The formula unit calculated on the basis of 24 O atoms is (K 0.15 Na 0.82 )(Na 1 . 97 Ca 0.03 )(Mn 2+ 1.66 Mg 0.34 - Mn 3+ 2.96 Fe 3+ 0.06 Zn 0.01 )(Si 7.99 Al 0.01 )O 2 and is close to the ideal end-member composition of NaNa 2 (Mn 2+ 2 Mn 3+ 3 )Si 8 O 22 O 2 . The crystal structure of ungarettiite was refined to an R index of ~ 1.5% using MoKα X-ray intensity data. Site-scattering refinement shows that the M1, M2, and M3 sites are occupied dominantly by Mn. The (M-O) distances are M1: 2.03; M2: 2.17; M3: 2.01 Å , compatible with the site occupancies M1~ M3 ~ Mn 3+ , M2 ~ Mn 2+ . All bonds to the O3 anion are very short, and a bond valence analysis indicates that the O3 site is occupied by a divalent anion: O 2- , as suggested by the overall electroneutrality requirement of the structural formula.

The crystal structure of minehillite from Franklin, New Jersey, (K,Na) 2 Ca 28 Zn 5 Al 4 Si 40 - O 112 (OH) 16 , was solved and refined in space group P3̅>c1, a = 9.777(2), c = 33.293(2) Å, to R = 0.022 for 1510 unique reflections collected from a twinned crystal. Two twinned portions are related to each other by a twofold axis coincident with the 3̅ axis, which can result in apparent hexagonal symmetry for minehillite. The structure consists of a stacked sequence of three types of layer units: (I) an infinite sheet of edge-sharing Ca-(O,OH) polyhedra, (2) a single sheet of SiO 4 tetrahedra connected in oval and pseudohexagonal­shaped six-membered rings, and (3) a complex slab built of SiO 4 tetrahedra and AlO 6 octahedra into which alkali elements and Zn are accommodated. The first two units are identical to those found in the reyerite structure. The third is analogous to the double layer in reyerite but with significant differences. The SiO 4 tetrahedra in the complex slab in minehillite form two sheets of isolated pseudohexagonal six-membered rings that are con­nected by AlO 6 octahedra centered on the threefold axis, which are unusual features among layered silicate minerals. An average of five Zn atoms per cell partially occupy a highly distorted tetrahedral site.

Artroeite, PbAlF 3 (OH) 2 , space group P1̄, a = 6.270(2), b = 6.821(3), c = 5.057(2) Å, α = 90.68(2), β = 107.69(2), γ= 104.46(2)°, V = 198.6(2) Å 3 , Z = 2, is a new mineral from the Grand Reef mine, near Klondyke, Graham County, Arizona. It occurs as colorless bladed crystals associated with quartz, fluorite, galena, anglesite, and an as yet undescribed mineral of composition PbCa 2 Al(F,OH) 9 . Artroeite has a Mohs hardness of about 2.5, a measured density of 5.36(2) g/cm 3 , and a calculated density of 5.47 g/cm 3 . It exhibits a perfect {100} cleavage and a good {010} cleavage. Optically it is biaxial (-) with α = 1.629(1), β = 1.682(2), and γ = 1.691(2). Dispersion is strong, r > ν. The six strongest powder diffraction lines are [d(I,hkl)]4.42 100 (1̄ 01), 3.221 40 (101), 2.595 70 (12̅1,021), 2.190 65 (201,012,030), 2.030 50 (022), 2.015 40 (2̅30) Å. The structure was solved by direct methods and refined to R = 0.022 using X-ray diffractometer data (1096 independent reflections). In the structure, edge-sharing dimers of AlF 3 (OH) 3 octahedra link together by bonds to Pb atoms to form approximately close- packed layers parallel to (101̄). The layers are linked to one another by one Pb-F bond and two H bonds per formula unit. Pb is coordinated to 6 F and 3 O atoms. Three F ligands are associated with the same Al octahedral face and correspond to much longer bonds. It is probable that the lone pair electrons are situated on that side of the Pb atom. The structure is compared with those of acuminite and tikhonenkovite.