Nancyrossite, ideally FeGeO6H5, is a new hydroxyperovskite from Tsumeb. It probably formed by the oxidation and partial dehydrogenation of stottite, FeGe(OH)6, with which it is associated intimately. The structure of nancyrossite has been determined in tetragonal space group P42/n: a = 7.37382(12) & Aring;, c = 7.29704(19) & Aring;, V = 396.764(16) & Aring;3, Z = 4, R1(all) = 0.034, wR2(all) = 0.051 and GoF = 1.057. Empirical formulae of two crystals have almost end-member compositions, Fe3+1.01Zn0.03Ge0.98O6H5 and Fe3+1.01Zn0.04Ge0.98O6H5. Structure determination indicates that 88% of the Fe is ferric. The chemical formula proposed here for nancyrossite recognises that although H atoms form OH groups, writing the formula as FeGeO(OH)5 implies that one of the six oxygen atoms is very underbonded, with a bond-valence sum of only similar to 1.2 valence units. As such, H in nancyrossite may have novel crystal chemistry. For example, the five H atoms may be distributed dynamically over the six O atoms, a phenomenon that would be averaged by X-ray diffraction, and so go undetected. Nancyrossite is the Ge-analogue of jeanbandyite. By analogy with nancyrossite, we propose revision of the ideal formula of jeanbandyite from FeSnO(OH)5 (Welch and Kampf, 2017) to FeSnO6H5.

This investigation investigates geologically old, ca. 370 Ma, metamict zirconolite from the Kovdor phoscorites and carbonatites in the Kola Alkaline Province. Mineral composition, crystallisation behaviour, and thermal expansion of the recrystallised samples were analysed using electron microprobe analysis, Raman spectroscopy, and in situ high-temperature powder X-ray diffraction (HTPXRD). The zirconolite crystals investigated are different in their morphology, internal texture, composition, alteration degree, and can be divided into four distinct groups. The zirconolite is a high Nb and Fe3+ variety (10.8-24.1 wt.% Nb2O5 and 7.9-9.0 wt.% Fe2O3), it is enriched in Th (up to 8.7 wt.% ThO2), Ta (up to 5.3 wt.% Ta2O5) and rare earth elements (up to 5.0 wt.% REE2O3). Raman spectroscopy confirmed that metamict zirconolite is anhydrous. The recrystallisation process of the metamict zirconolite is complex, as detected by HTPXRD. A fluorite-type phase starts to crystallise at 420 degrees C. The formation of a pyrochlore phase can be identified at 750 degrees C. The major phases detected in the sample after the recrystallisation are: zirconolite-3T (53 wt.%), srilankite (25 wt.%), pyrochlore (15 wt.%), baddeleyite (5 wt.%) and zircon (3 wt.%). The average coefficients of thermal expansion (CTE) values in the temperature range 25-1200 degrees C are as follows: ${{\bar \alpha }}$a = ${{\bar \alpha }}$b = ${{\bar \alpha }}$11 = ${{\bar \alpha }}$22 = 8.9510-6 deg-1. Similarly, the thermal expansion along the c-axis yields a similar value: ${{\bar \alpha }}$a = ${{\bar \alpha }}$b = 8.9310-6 deg-1, indicating an almost isotropic thermal expansion of zirconolite-3T. The lower CTE value compared to a pure synthetic zirconolite observed for zirconolite-3T might be attributed to the complex chemistry and polyphase nature of the material investigated.This investigation investigates geologically old, ca. 370 Ma, metamict zirconolite from the Kovdor phoscorites and carbonatites in the Kola Alkaline Province. Mineral composition, crystallisation behaviour, and thermal expansion of the recrystallised samples were analysed using electron microprobe analysis, Raman spectroscopy, and in situ high-temperature powder X-ray diffraction (HTPXRD). The zirconolite crystals investigated are different in their morphology, internal texture, composition, alteration degree, and can be divided into four distinct groups. The zirconolite is a high Nb and Fe3+ variety (10.8-24.1 wt.% Nb2O5 and 7.9-9.0 wt.% Fe2O3), it is enriched in Th (up to 8.7 wt.% ThO2), Ta (up to 5.3 wt.% Ta2O5) and rare earth elements (up to 5.0 wt.% REE2O3). Raman spectroscopy confirmed that metamict zirconolite is anhydrous. The recrystallisation process of the metamict zirconolite is complex, as detected by HTPXRD. A fluorite-type phase starts to crystallise at 420 degrees C. The formation of a pyrochlore phase can be identified at 750 degrees C. The major phases detected in the sample after the recrystallisation are: zirconolite-3T (53 wt.%), srilankite (25 wt.%), pyrochlore (15 wt.%), baddeleyite (5 wt.%) and zircon (3 wt.%). The average coefficients of thermal expansion (CTE) values in the temperature range 25-1200 degrees C are as follows: ${{\bar \alpha }}$a = ${{\bar \alpha }}$b = ${{\bar \alpha }}$11 = ${{\bar \alpha }}$22 = 8.9510-6 deg-1. Similarly, the thermal expansion along the c-axis yields a similar value: ${{\bar \alpha }}$a = ${{\bar \alpha }}$b = 8.9310-6 deg-1, indicating an almost isotropic thermal expansion of zirconolite-3T. The lower CTE value compared to a pure synthetic zirconolite observed for zirconolite-3T might be attributed to the complex chemistry and polyphase nature of the material investigated.

The Cornubian Batholith (SW England) is an archetypal Variscan rare metal granite with potential for Li-mica mineralization. We present a petrographic, trace element and multivariate statistical study of micas from the Cornubian Batholith granite series and related hydrothermally altered units to assess the role of magmatic vs subsolidus processes and of fluxing elements (F and B) on the Li cycle during the evolution of the system. The mica types are as follows: (1) magmatic, which include Fe-biotite, protolithionite I and phengite-muscovite from the most primitive granites, and zinnwaldite I from more fractionated lithologies; (2) subsolidus, which encompass high-temperature autometasomatic Li-micas and low-temperature hydrothermal muscovite-phengite. Autometasomatic species include protolithionite II, zinnwaldite II and lepidolite, which were observed in the most fractionated and hydrothermally altered units, and occur as replacements of magmatic micas. Low-temperature hydrothermal Li-poor micas formed via alteration of magmatic and autometasomatic micas or as replacement of feldspars, and albeit occur in all studied lithologies they are best represented by the granite facies enriched in metasomatic tourmaline. The evolution of micas follows two major trends underlining a coupling and decoupling between the Li(F) and B fluxes. These include as follows: (1) a Li(F)-progressive trend explaining the formation of protolithionite I and zinnwaldite I, which fractionate Li along with Cs, Nb and Sn during the late-magmatic stages of crystallization, and of zinnwaldite II and lepidolite forming from the re-equilibration of primary micas with high-temperature Li-B-W-Tl-Cs-Mn-W-rich autometasomatic fluids; (2) a Li(F)-retrogressive trend explaining the low-temperature hydrothermal muscovitization, which represents the main Li depletion process. Trace element geochemistry and paragenesis of late muscovite-phengite support that muscovitization is a district-scale process that affected the upper parts of the granite cupolas through acidic and B(Fe-Sn)-saturated hydrothermal fluids associated with metasomatic tourmalinization, which were mixed with a low Eh meteoric component.

Almost all igneous rocks are composed of silicate minerals; carbonatites are the main exception to this rule. They form only a minor proportion of the continental crust but are of fundamental scientific and economic importance. These rocks, originally described as a limestone (Kaiserstuhl, Germany) or magmatic limestone (Alnö, Sweden), were recognized in 1921 by W. C. Brøgger as a distinct magmatic rock type under the name “Karbonatite” in the Fen complex (Norway). Extensive field mapping in Africa, and particularly studies within the Chilwa Alkaline province (Malawi), have led to the discovery of diverse intrusive and extrusive carbonatites. The latter, including the Fort Portal volcanic field in Uganda, Rufunsa Province in Zambia, and the Oldoinyo Lengai volcano in Tanzania (the only volcano ever to have been seen to erupt carbonatites), have been exceptionally important in the recognition of carbonatites as truly magmatic rocks. The possibility of the existence of carbonate melts has been confirmed and shown by experimental studies of diverse carbonate systems with added volatile components (H2O, F) and alkali elements (Na, K). The study of the Oldoinyo Lengai gregoryite-nyerereite carbonatites, which are mineralogically and compositionally different from all known carbonatites worldwide, has led to long-lasting discussions about the origin of carbonatites. This includes composition of primary/parental carbonate melt, derivation of carbonatites by either liquid immiscibility or fractional crystallization, carbonatite evolution and especially, the possible genetic relationships between alkali-rich and alkali-poor carbonatites. The rapid alteration of Oldoinyo Lengai carbonatites and their transformation to calcite carbonatite-like rocks has been proposed as the explanation for the absence of alkali-rich carbonatites in the geological past. Detailed mineralogical studies have shown that the occurrence of nyerereite is not restricted to Oldoinyo Lengai and that this mineral is now known to occur in other carbonatites (e.g., Guly, Kovdor, Oka, Kerimasi), alkaline rocks, and kimberlites (and even in diamond). This would suggest compositionally different mantle-derived melts enriched in alkali elements. In addition, carbonatite tephra has an important role in the preservation of some key paleontologic and anthropologic localities in East Africa. Despite these important discoveries, several problems related to carbonatite petrogenesis are not resolved yet; future work is required, and carbonatites within Africa with its key localities may help to achieve them.

Exposure of the Hubble Space Telescope to space in low Earth orbit resulted in numerous hypervelocity impacts by cosmic dust (micrometeoroids) and anthropogenic particles (orbital debris) on the solar arrays and the radiator shield of the Wide Field and Planetary Camera 2, both subsequently returned to Earth. Solar cells preserve residues from smaller cosmic dust (and orbital debris) but give less reliable information from larger particles. Here, we present images and analyses from electron, ion and X-ray fluorescence microscopes for larger impact features (millimetre- to centimetre-scale) on the radiator shield. Validated by laboratory experiments, these allow interpretation of composition, probable origin and likely dimensions of the larger impactors. The majority (~90%) of impacts by grains greater than 50 μm in size were made by micrometeoroids, dominated by magnesium- and iron-rich silicates and iron sulfides, metallic iron-nickel and chromium-rich spinel similar to that in ordinary chondrite meteorites of asteroid origin. Our re-evaluation of the largest impact features shows substantially fewer large orbital debris impacts than reported by earlier authors. Mismatch to the NASA ORDEM and ESA MASTER models of particle populations in orbit may be partly due to model overestimation of orbital debris flux and underestimation of larger micrometeoroid numbers. This article is part of the theme issue ‘Dust in the Solar System and beyond’.

Mikecoxite, ideally (CHg4)OCl2, is the first mercury-oxide-chloride-carbide containing a C4- anion coordinated by four Hg atoms (a permercurated methane derivative) to be described as a mineral species. It was found at the McDermitt open-pit mine on the eastern margin of the McDermitt Caldera, Humboldt County, Nevada, U.S.A. It is monoclinic, space group P21/n, Z = 4; a = 10.164(5), b = 10.490(4), c = 6.547(3) A, V 698.0(5) A(3). Chemical analysis by electron microprobe gave Hg 86.38, Cl 11.58, Br 0.46, C 1.81, sum = 100.23 wt%, and O was detected but the signal was too weak for quantitative chemical analysis. The empirical formula, calculated on the basis of Hg + Cl + Br = 6 apfu, is (C1.19Hg3.39)(Cl2.57Br0.05)(sigma 2.62), and the ideal formula based on the chemical analysis and the crystal structure is (CHg4)OCl2. The seven strongest lines in the X-ray powder diffraction pattern are [d (A), I, (hkl)]: 2.884, 100, (230); 2.989, 81, (301, 301, 112, 112, 131, 131); 2.673, 79, (122, 122, 212, 212); 1.7443, 40, (060, 432, 432); 5.49, 34, (101, 101); 4.65, 32, (120); 2.300, 30, (312, 312). The Raman spectrum shows three bands at 638, 675, and 704 cm(-1), well above the range characteristic of NHg4 stretching vibrations between 540 and 580 cm(-1), that are assigned to CHg4 stretching vibrations. Mikecoxite forms intergrowths of bladed crystals up to 100 mu m long that occur on granular quartz or in vugs associated with kleinite. It is black with a submetallic to metallic luster and strong specular reflections and does not fluoresce under short- or long-wave ultraviolet light. Neither cleavage nor parting were observed, and the calculated density is 8.58 g/cm(3). In the crystal structure of mikecoxite, ( C (4 -) H g (4) (2 +) ) groups link through O2- ions to form three-membered rings that polymerize into corrugated [CHg4OCl](+) layers with near-linear C4--Hg2+-O and C4--Hg2+-Cl linkages. The layers link in the third direction directly via weak Hg2+-O2- and Hg2+-Cl- bonds to adjacent layers and also indirectly via interlayer Cl-. A bond-valence parameter has been derived for (Hg2+-C4-) bonds: R-o = 2.073 &A, b = 0.37, which gives bond-valence sums at the C4- ions in accord with the valence-sum rule. The source of carbon for mikecoxite in the volcanic high-desert environment of the type locality seems to be methane, with the reaction catalyzed by microbiota through full mercuration of carbon atoms, beyond the first stage that produces the volatile and highly mobile methylmercury, [CH3Hg](+), a potent neurotoxin that accumulates in marine food chains. Both the mineral and the mineral name have been approved by the Commission on New Minerals, Nomenclature and Classification of the International Mineralogical Association (IMA 2021-060). The mineral is named after Michael F. Cox (b. 1958), a founding member of the New Almaden Quicksilver County Park Association (NAQCPA) who was responsible for characterizing and remediating environmental mercury on-site and who recovered the rock containing the new mineral.

We searched late Miocene sedimentary rocks in an attempt to recover fossil micrometeorites derived from the Veritas asteroid family. This study was motivated by the previous identification of a pronounced He-3 peak (4-5x above background) within marine sediments with ages between similar to 8.5-6.9 Ma ago (Montanari et al., 2017. GSA Bulletin, 129:1357-1376). We processed 118.9 kg of sediment from the Monte dei Corvi beach section (Italy), the global type-section for the Tortonian epoch (11.6-7.2 Ma). Samples were collected both before and within the He-3 peak. Although a small number of iron-rich (I-type) fossil micrometeorites were recovered from each horizon studied (N-total = 20), there is no clear difference between the pre- and intra- He-3 peak samples. All micrometeorites are compositionally similar, and three out of five horizons yielded similar abundances and particle sizes. Micrometeorites extracted from sediments at the base of the He-3 peak were exclusively small (phi <75 mu m), while micrometeorites extracted from sediments near the highest He-3 values were relatively large (phi <270 mu m). The recovered fossil micrometeorites are interpreted as samples of the background dust flux derived from metal-bearing chondritic asteroids. The presence of a He-3 signature combined with the absence of fossil micrometeorites or extraterrestrial spinels (Boschi et al., 2019, Spec. Pap. Geol. Soc. Am. 542:383-391) unambiguously related to the Veritas event suggests that the Veritas family is composed of highly friable materials that rarely survive on the sea floor to become preserved in the geological record. Our data supports the existing hypothesis that the Veritas asteroid family is an aqueously altered carbonaceous chondrite parent body, one that contains minimal native metal grains or refractory Cr-spinels. The low yield of fossil micrometeorites at Monte dei Corvi is attributed to loss of particles by dissolution whilst they resided on the sea floor but also due to high sedimentation rates leading to dilution of the extraterrestrial dust flux at this site. As with other fossil micrometeorite collections (e.g. Cretaceous chalk [Suttle and Genge, EPSL, 476:132-142]) the I-type spherules have been altered since deposition. In most particles, both magnetite and wustite remain intact but have been affected by solid state geochemical exchange, characterised by partial leaching of Ni, Co and Cr and implantation of Mn, Mg, Si and Al. In some particles Mn concentrations reach up to 16.6 wt%. Conversely, in some micrometeorites wustite has been partially dissolved, or even replaced by calcite or ankerite. Finally, we observe evidence for wustite recrystallisation, forming a second generation of magnetite. This process is suggested to occur by oxidation during residence on the seafloor and has implications for the use of fossil I-type micrometeorites as a potential proxy for probing Earth's upper atmospheric composition (oxidative capacity) in the geological past. However, solutions to the limitations of post-depositional recrystallisation are suggested. Fossil I-type spherules remain a potential tool for palaeo-climatic studies.

Geological characteristics, trace element geochemistry, and Re-Os dating of molybdenite are presented and discussed for seven Cu and Mo porphyry, one Cu skarn-porphyry and two Mo and Mo-W greisen deposits. Re contents in molybdenite are discussed versus ore mineralogy, molybdenite polytypes, formation temperature, and whole rock chemistry of parental rocks. They are also evaluated in the light of the geotectonic settings of these deposits and tectonic history of the Urals.Re-Os dating of molybdenite evidences that the porphyry deposits studied are confined to four volcanic arc terranes of the Urals and are linked to four subduction events: Tomino and Birgilda (ca. 430 Ma) and Zeleny Dol (ca. 420 Ma) porphyry deposits linked to the East Uralian volcanic arc; Voznesenskoe porphyry (ca. 395 Ma) and Verkhneuralskoe porphyry (ca. 365 Ma) deposits linked to the Magnitogorsk arc and its collision to the East European continent respectively; Mikheevskoe porphyry and Tarutino skarn-porphyry deposits (ca. 360 Ma) linked to subduction under the East Uralian microcontinent; Benkala porphyry deposit (ca. 335 Ma) linked to the activity of the Valerianovka Andean-type arc. The Talitsa Mo porphyry deposit (ca. 300 Ma) as well as YuzhnoShameiskoe (ca. 280 Ma) and Koklanovskoe (ca. 255 Ma) greisen Mo-W deposits are linked to collision between East European and Kazakh continents.Electron microprobe analysis and laser ablation inductively coupled mass-spectrometry of molybdenite demonstrate an extremely irregular distribution of all trace elements on both grain-scale and deposit-scale levels. Most trace elements (Fe, Co, Cu, Zn, As, Se, Ag, Sb, Te, Pb, Bi, etc.) form mineral inclusions within molybdenite while Re and W are commonly incorporated into molybdenite lattice. Re contents and distribution in molybdenite are indicative for tectonic setting of the deposits. Molybdenite from deposits formed within intra-oceanic arcs (Mikheevskoe, Tomino, and Voznesenskoe porphyry Cu deposits) is featured by high Re contents (1000-5000 ppm) and low W contents (<10 ppm). Molybdenite from Verkhneuralskoe (arc-continent collision stage) and Benkala (Andean-type arc) deposits is featured by lower Re contents (mainly 400-900 ppm) and higher W contents (30-120 ppm). Molybdenite from Talitsa Mo-porphyry, and Yuzhno-Shameiskoe and Koklanovskoe greisen deposits (continent-continent collision) is featured by the lowest Re contents (<10 ppm to 370 ppm) and the highest W contents (10 to 150 ppm). Geometrical mean Re content in molybdenite for all ten deposits studied has positive correlation with average Cu/Mo ratio in ore implying mass balance was one of the key controls of Re incorporation in molybdenite. In addition, Re content has a negative correlation with SiO2 and total REE contents in ore-bearing intrusive rocks as well as with the absolute age of the deposits which correlates with the general evolution of porphyry- and greisen-type deposits of the Urals. For subduction-related porphyrydeposits magma composition was generated with strong influence of the shallow subduction component. Fluid -mobile elements and various degrees of partial melting of the mantle wedge strongly depend on Re contents. Collision tectonic settings are marked by low Re concentrations in molybdenite.

Determination of alkali elements is important to Earth scientists, yet suitable and reliable microanalytical reference materials are lacking. This paper proposes a new albite reference material and evaluates the potential for future K‐feldspar reference materials. The proposed Piz Beverin albite reference material from Switzerland yields a homogeneous composition at the centimetre‐ to micrometre‐scale for Si, Al and Na with < 2000 μg g ‐1 total trace elements (mostly heterogeneously distributed Ca, K and Sr). EPMA and LA‐ICP‐MS measurements confirm a composition of 99.5(2)% albite component, which is supported further by bulk XRF measurements. A round robin evaluation involving nine independent EPMA laboratories confirms its composition and homogeneity for Si, Al and Na. In addition, a set of five distinct clear K‐feldspar samples was evaluated as possible reference materials. The first two crystals of adular and orthoclase yield unacceptable inhomogeneities with > 2% relative local variations of Na, K and Ba contents. The three other investigated sets of K‐feldspar crystals are yellow sanidine crystals from Itrongay (Madagascar). Despite distinct compositions, EPMA confirms they are each homogeneous at the centimetre to micrometre scale for Si, Al and K and have no apparent inclusions; further investigation to find larger amounts of these materials is therefore justified.

Synthetic and naturally occurring forms of tricopper orthotellurate, CuII3TeVIO6 (the mineral mcalpineite) have been investigated by 3D electron diffraction (3D ED), X-ray powder diffraction (XRPD), Raman and infrared (IR) spectroscopic measurements. As a result of the diffraction analyses, CuII3TeVIO6 is shown to occur in two polytypes. The higher-symmetric CuII3TeVIO6-1C polytype is cubic, space group Ia3, with a = 9.537 (1) Å and V = 867.4 (3) Å3 as reported in previous studies. The 1C polytype is a well characterized structure consisting of alternating layers of CuIIO6 octahedra and both CuIIO6 and TeVIO6 octahedra in a patchwork arrangement. The structure of the lower-symmetric orthorhombic CuII3TeVIO6-2O polytype was determined for the first time in this study by 3D ED and verified by Rietveld refinement. The 2O polytype crystallizes in space group Pcca, with a = 9.745 (3) Å, b = 9.749 (2) Å, c = 9.771 (2) Å and V = 928.3 (4) Å3. High-precision XRPD data were also collected on CuII3TeVIO6-2O to verify the lower-symmetric structure by performing a Rietveld refinement. The resultant structure is identical to that determined by 3D ED, with unit-cell parameters a = 9.56157 (19) Å, b = 9.55853 (11) Å, c = 9.62891 (15) Å and V = 880.03 (2) Å3. The lower symmetry of the 2O polytype is a consequence of a different cation ordering arrangement, which involves the movement of every second CuIIO6 and TeVIO6 octahedral layer by (1/4, 1/4, 0), leading to an offset of TeVIO6 and CuIIO6 octahedra in every second layer giving an ABAB* stacking arrangement. Syntheses of CuII3TeVIO6 showed that low-temperature (473 K) hydrothermal conditions generally produce the 2O polytype. XRPD measurements in combination with Raman spectroscopic analysis showed that most natural mcalpineite is the orthorhombic 2O polytype. Both XRPD and Raman spectroscopy measurements may be used to differentiate between the two polytypes of CuII3TeVIO6. In Raman spectroscopy, CuII3TeVIO6-1C has a single strong band around 730 cm-1, whereas CuII3TeVIO6-2O shows a broad double maximum with bands centred around 692 and 742 cm-1.

The crystal structure of montanite has been determined using single-crystal X-ray diffraction on a synthetic sample, supported by powder X-ray diffraction (PXRD), electron microprobe analysis (EPMA) and thermogravimetric analyses (TGA). Montanite was first described in 1868 as Bi2TeO6·nH2O (n = 1 or 2). The determination of the crystal structure of synthetic montanite (refined composition Bi2TeO6·0.22H2O) has led to the reassignment of the formula to Bi2TeO6·nH2O where 0 ≤ n ≤  $${\raise0.5ex\hbox{$\scriptstyle 2$} \kern-0.1em/\kern-0.15em \lower0.25ex\hbox{$\scriptstyle 3$}}$$ rather than the commonly reported Bi2TeO6·2H2O. This change has been accepted by the IMA–CNMNC, Proposal 22-A. The PXRD pattern simulated from the crystal structure of synthetic montanite is a satisfactory match for PXRD scans collected on both historical and recent natural samples, showing their equivalence. Two specimens attributed to the original discoverer of montanite (Frederick A. Genth) from the cotype localities (Highland Mining District, Montana and David Beck’s mine, North Carolina, USA) have been designated as neotypes. Montanite crystallises in space group P $$\overline{6 }$$ , with the unit-cell parameters a = 9.1195(14) Å, c = 5.5694(8) Å, V = 401.13(14) Å3, and three formula units in the unit cell. The crystal structure of montanite is formed from a framework of BiOn and TeO6 polyhedra. Half of the Bi3+ and all of the Te6+ cations are coordinated by six oxygen atoms in trigonal-prismatic arrangements (the first example of this configuration reported for Te6+), while the remaining Bi3+ cations are coordinated by seven O sites. The H2O groups in montanite are structurally incorporated into the network of cavities formed by the three-dimensional framework, with other cavity space occupied by the stereoactive 6s2 lone pair of Bi3+ cations. While evidence for a supercell was observed in synthetic montanite, the subcell refinement of montanite adequately indexes all reflections in the PXRD patterns observed in all natural montanite samples analysed in this study, verifying the identity of montanite as a mineral.

The crystal structure of montanite has been determined using single-crystal X-ray diffraction on a synthetic sample, supported by powder X-ray diffraction (PXRD), electron microprobe analysis (EPMA) and thermogravimetric analyses (TGA). Montanite was first described in 1868 as Bi2TeO6 center dot nH(2)O (n = 1 or 2). The determination of the crystal structure of synthetic montanite (refined composition Bi2TeO6 center dot 0.22H(2)O) has led to the reassignment of the formula to Bi2TeO6 center dot nH(2)O where 0 <= n <= 2/3) rather than the commonly reported Bi2TeO6 center dot 2H(2)O. This change has been accepted by the IMA-CNMNC, Proposal 22-A. The PXRD pattern simulated from the crystal structure of synthetic montanite is a satisfactory match for PXRD scans collected on both historical and recent natural samples, showing their equivalence. Two specimens attributed to the original discoverer of montanite (Frederick A. Genth) from the cotype localities (Highland Mining District, Montana and David Beck's mine, North Carolina, USA) have been designated as neotypes. Montanite crystallises in space group P (6) over bar, with the unit-cell parameters a = 9.1195(14) angstrom, c = 5.5694(8) angstrom, V = 401.13(14) angstrom 3, and three formula units in the unit cell. The crystal structure of montanite is formed from a framework of BiOn and TeO6 polyhedra. Half of the Bi3+ and all of the Te6+ cations are coordinated by six oxygen atoms in trigonal-prismatic arrangements (the first example of this configuration reported for Te6+), while the remaining Bi3+ cations are coordinated by seven O sites. The H2O groups in montanite are structurally incorporated into the network of cavities formed by the three-dimensional framework, with other cavity space occupied by the stereoactive 6s(2) lone pair of Bi3+ cations. While evidence for a supercell was observed in synthetic montanite, the subcell refinement of montanite adequately indexes all reflections in the PXRD patterns observed in all natural montanite samples analysed in this study, verifying the identity of montanite as a mineral.

AbstractBridgesite-(Ce), (IMA2019-034), was discovered at Tynebottom Mine, Cumbria, UK. It occurs as thin (1–2 μm) translucent blue crystals with a lath-like to acicular habit, aggregated into thin crusts and is associated mainly with brochantite, malachite, serpierite, devilline, gypsum, aragonite, jarosite, pyrite, lanthanite-(Ce) and undifferentiated iron oxyhydroxides, it is often intergrown with these other minerals. The lustre, hardness, cleavage and parting could not be determined, nor could density be measured due to crystal size. It has a pale blue streak and is brittle with a splintery fracture. Bridgesite-(Ce) is biaxial (–), shows no pleochroism and has refractive indices (white light): α = 1.526(2), β = 1.564(2), γ = 1.572(2) and 2V(calc) = 48.3°. The empirical formula calculated on the basis of 44 negative charges is Ca0.86REEΣ1.99Al0.07Cu5.95(SO4)3.99(SiO4)0.05(PO4)0.02(OH)11.52⋅8H2O. The idealised formula is CaCe2Cu6(SO4)4(OH)12⋅8H2O, requiring (wt.%): 3.91 CaO, 22.89 Ce2O3, 33.28 CuO, 22.33 SO3 and 17.59 H2O. Bridgesite-(Ce) is monoclinic, space group C2/m, a = 24.801(5), b = 6.3520(13), c = 11.245(2) Å, β = 114.51(3)°, V = 1611.9(6) Å and Z = 2. The five most intense X-ray diffraction peaks in the measured pattern are [d in Å (I, %) (hkl)]: 11.3 (100) (200), 6.391 (15) (201), 2.770 (8) (420), 3.194 (6) (402) and 4.858 (5) (310). The crystal structure was solved using single crystal data and refined to an R1 index of 5.86%. Bridgesite-(Ce) contains three distinct Cu sites containing Cu2+, two are coordinated octahedrally and one is square pyramidal. The octahedra form chains through edge sharing parallel to the b-axis which are linked by the square pyramid to form sheets oriented parallel to {100}. Sulfate tetrahedra decorate the sheets which are held together by interstitial REE3+, Ca2+ and hydrogen bonding. The structure is unique. Despite apparent similarity in chemical formula, bridgesite-(Ce) is not closely related to any other natural Cu-sulfate mineral. An FTIR absorption spectra is presented for reference purposes.

AbstractThe crystal structure of montanite has been determined using single-crystal X-ray diffraction on a synthetic sample, supported by powder X-ray diffraction (PXRD), electron microprobe analysis (EPMA) and thermogravimetric analyses (TGA). Montanite was first described in 1868 as Bi2TeO6·nH2O (n = 1 or 2). The determination of the crystal structure of synthetic montanite (refined composition Bi2TeO6·0.22H2O) has led to the reassignment of the formula to Bi2TeO6·nH2O where 0 ≤ n ≤ $${\raise0.5ex\hbox{$\scriptstyle 2$} \kern-0.1em/\kern-0.15em \lower0.25ex\hbox{$\scriptstyle 3$}}$$ 2 / 3 rather than the commonly reported Bi2TeO6·2H2O. This change has been accepted by the IMA–CNMNC, Proposal 22-A. The PXRD pattern simulated from the crystal structure of synthetic montanite is a satisfactory match for PXRD scans collected on both historical and recent natural samples, showing their equivalence. Two specimens attributed to the original discoverer of montanite (Frederick A. Genth) from the cotype localities (Highland Mining District, Montana and David Beck’s mine, North Carolina, USA) have been designated as neotypes. Montanite crystallises in space group P$$\overline{6 }$$ 6 ¯ , with the unit-cell parameters a = 9.1195(14) Å, c = 5.5694(8) Å, V = 401.13(14) Å3, and three formula units in the unit cell. The crystal structure of montanite is formed from a framework of BiOn and TeO6 polyhedra. Half of the Bi3+ and all of the Te6+ cations are coordinated by six oxygen atoms in trigonal-prismatic arrangements (the first example of this configuration reported for Te6+), while the remaining Bi3+ cations are coordinated by seven O sites. The H2O groups in montanite are structurally incorporated into the network of cavities formed by the three-dimensional framework, with other cavity space occupied by the stereoactive 6s2 lone pair of Bi3+ cations. While evidence for a supercell was observed in synthetic montanite, the subcell refinement of montanite adequately indexes all reflections in the PXRD patterns observed in all natural montanite samples analysed in this study, verifying the identity of montanite as a mineral.

AbstractScroll-like crystals of molybdenite, 2–5 mm in size, were found in phengite rock from the outer contact of the granular quartz vein of the Kyshtym quartz deposit. Platy and partly scrolled molybdenite occur in the same phengite rock from the outer contact of the quartz–feldspar pegmatite of the Slyudyanogorsk mica deposit. Both occurrences are located in the Ufaley metamorphic block in the South Urals. Scroll-like molybdenite crystals can associate with platy and partly twisted crystals in the same samples. The chemical composition of molybdenite was studied by inductively coupled plasma mass spectrometry (ICP-MS) and electron probe microanalysis (EPMA). Polytypes of molybdenite were identified with electron back-scattered diffraction (EBSD) and X-ray diffraction (XRD). Both scroll-like and platy molybdenite crystals are only represented by the 3R polytype, are enriched in Re up to 1 wt.% and contain no other significant impurities. Scroll-like molybdenite is twisted mainly around the crystallographic axis X. Twinning with a rotation of 60 degrees around the Z crystallographic axis is fixed in the plane (ab). The most probable origin of scroll molybdenites is the consequent growth of molybdenite around nucleation centres, which are commonly represented by mica crystals. The formation of the 3R polytype is caused by the difference in dimension of the layers enriched and depleted in rhenium.

The crystal structure of montanite has been determined using single-crystal X-ray diffraction on a synthetic sample, supported by powder X-ray diffraction (PXRD), electron microprobe analysis (EPMA) and thermogravimetric analyses (TGA). Montanite was first described in 1868 as Bi 2 TeO 6 · n H 2 O ( n = 1 or 2). The determination of the crystal structure of synthetic montanite (refined composition Bi 2 TeO 6 ·0.22H 2 O) has led to the reassignment of the formula to Bi 2 TeO 6 · n H 2 O where 0 ≤ n ≤ 0.5ex2-0.1em/-0.15em 0.25ex3 rather than the commonly reported Bi 2 TeO 6 ·2H 2 O. This change has been accepted by the IMA–CNMNC, Proposal 22-A. The PXRD pattern simulated from the crystal structure of synthetic montanite is a satisfactory match for PXRD scans collected on both historical and recent natural samples, showing their equivalence. Two specimens attributed to the original discoverer of montanite (Frederick A. Genth) from the cotype localities (Highland Mining District, Montana and David Beck’s mine, North Carolina, USA) have been designated as neotypes. Montanite crystallises in space group P 6 , with the unit-cell parameters a = 9.1195(14) Å, c = 5.5694(8) Å, V = 401.13(14) Å 3 , and three formula units in the unit cell. The crystal structure of montanite is formed from a framework of BiO n and TeO 6 polyhedra. Half of the Bi 3+ and all of the Te 6+ cations are coordinated by six oxygen atoms in trigonal-prismatic arrangements (the first example of this configuration reported for Te 6+ ), while the remaining Bi 3+ cations are coordinated by seven O sites. The H 2 O groups in montanite are structurally incorporated into the network of cavities formed by the three-dimensional framework, with other cavity space occupied by the stereoactive 6 s 2 lone pair of Bi 3+ cations. While evidence for a supercell was observed in synthetic montanite, the subcell refinement of montanite adequately indexes all reflections in the PXRD patterns observed in all natural montanite samples analysed in this study, verifying the identity of montanite as a mineral.

The crystal structure of montanite has been determined using single-crystal X-ray diffraction on a synthetic sample, supported by powder X-ray diffraction (PXRD), electron microprobe analysis (EPMA) and thermogravimetric analyses (TGA). Montanite was first described in 1868 as Bi 2 TeO 6 · n H 2 O ( n = 1 or 2). The determination of the crystal structure of synthetic montanite (refined composition Bi 2 TeO 6 ·0.22H 2 O) has led to the reassignment of the formula to Bi 2 TeO 6 · n H 2 O where 0 ≤ n ≤ $${\raise0.5ex\hbox{$\scriptstyle 2$} \kern-0.1em/\kern-0.15em \lower0.25ex\hbox{$\scriptstyle 3$}}$$ 2 / 3 rather than the commonly reported Bi 2 TeO 6 ·2H 2 O. This change has been accepted by the IMA–CNMNC, Proposal 22-A. The PXRD pattern simulated from the crystal structure of synthetic montanite is a satisfactory match for PXRD scans collected on both historical and recent natural samples, showing their equivalence. Two specimens attributed to the original discoverer of montanite (Frederick A. Genth) from the cotype localities (Highland Mining District, Montana and David Beck’s mine, North Carolina, USA) have been designated as neotypes. Montanite crystallises in space group P $$\overline{6 }$$ 6 ¯ , with the unit-cell parameters a = 9.1195(14) Å, c = 5.5694(8) Å, V = 401.13(14) Å 3 , and three formula units in the unit cell. The crystal structure of montanite is formed from a framework of BiO n and TeO 6 polyhedra. Half of the Bi 3+ and all of the Te 6+ cations are coordinated by six oxygen atoms in trigonal-prismatic arrangements (the first example of this configuration reported for Te 6+ ), while the remaining Bi 3+ cations are coordinated by seven O sites. The H 2 O groups in montanite are structurally incorporated into the network of cavities formed by the three-dimensional framework, with other cavity space occupied by the stereoactive 6 s 2 lone pair of Bi 3+ cations. While evidence for a supercell was observed in synthetic montanite, the subcell refinement of montanite adequately indexes all reflections in the PXRD patterns observed in all natural montanite samples analysed in this study, verifying the identity of montanite as a mineral.

Igneous and hydrothermal mineral chemistry is commonly utilized in the exploration for porphyry deposits. Despite its occurrence in igneous and alteration mineral assemblages, routine analysis of titanite in mineral exploration is rare. We present whole-rock and titanite chemistry from the giant Cobre Panama porphyry Cu district to distinguish between primary igneous and replacement titanite. We show that igneous titanite is characterized by high Fe/Al ratios, REE, Zr and Mn concentrations, whereas F and Sc are enriched in replacement titanite. These distinctions can be explained by differences in mineral-melt partitioning, mobility in hydrothermal fluids, or by inheritance from precursor igneous minerals. Furthermore, REE patterns of igneous titanite reflect whole-rock chemistry, with discrepancies in MREEs possibly indicating that significant fractionation of amphibole has occurred prior to titanite saturation.

The Laetoli area in northern Tanzania is an important palaeo-anthropological site, where the oldest footprints of Australopithecus afarensis reside. Aeolian tuffs are the major rock type at Laetoli and they are divided into Lower and Upper Laetolil Beds that were deposited at an interval of 4.36 and 3.63 million years. The Upper Laetolil Beds contain eight layers of air-fall tuffs known as marker tuffs. The Australopithecus afarensis footprints are observed on the surface of the white tuff, which is a part of the Upper Laetolil marker tuff 7, also known as the “Footprint Tuff.” The interpolated age of the marker tuff 7 is 3.66 million years. Two mineral assemblages are distinguished in the Upper Laetolil marker tuffs. The first assemblage consists of primary tuff minerals and includes clinopyroxene (diopside, augite, aegirine-augite), nepheline, melilite (åkermanite and alumoåkermanite), garnet (andradite and schorlomite), magnetite, and others. The second mineral assemblage consists of secondary minerals, montmorillonite, calcite, and phillipsite. They were formed during replacement of the primary minerals, volcanic glass, and ash cementation. Thermodynamic calculations show that the major primary tuff minerals (melilite and nepheline) are stable at variable sodium activity and pH values. Replacement of melilite and nepheline by montmorillonite is caused by a decrease of sodium activity in slightly alkaline, neutral and acidic conditions (рН < 10). Montmorillonite is not present in the altered nephelinitic tuff of the Sadiman volcano (which is considered as a source of the Laetolil Beds) where kaolinite is the major secondary mineral. This is explained by the difference in H2O fugacity with higher lgfH2O values in Sadiman and lower values in Laetoli. Relationships between primary and secondary tuffs minerals on the lgaHCO3 vs pH plot suggest mineral transformation within the Laetolil Beds in slightly acid and neutral con- ditions (рН = 5–7) compared with more alkaline conditions at Sadiman (рН > 10).

Pegmatite fields within granite plutons are commonly considered to have formed from residual melts of their host. This is not always true as demonstrated by the Tysfjord granite gneiss and its two groups of pegmatites. The Tysfjord granite gneiss, exposed in a tectonic window of the Caledonides of northern Norway, is part of the transscandinavian igneous belt (TIB) that includes several phases of granitic magmatism. In the northern Hamarøy area (Drag-Finnøy), where most rare-element pegmatites occur, Paleoproterozoic and metamorphosed Group 1 allanite–(Ce)–fluorite metapegmatites have similar bulk rock chemical composition as the TIB granite gneiss rocks, indicating that these pegmatites are residual melts. Group 1 metapegmatites, which are up to 400 m in size, are among the largest known intra-plutonic pegmatites with Nb–Y–F (NYF) signature. The formation of these unusually large granite-hosted NYF pegmatites may have been facilitated by the overall high F content of TIB granite gneisses. Undeformed Group 2 amazonite–tourmaline pegmatites yield columbite and zircon U–Pb ages in the range 400–379 Ma. These pegmatites are interpreted to be anatectic melts that formed from the partial melting of Tysfjord granite gneiss. Group 2 pegmatites, including those from Træna Island and the Sjona tectonic window (400 and 414 Ma), formed during late Caledonian ductile shearing and incipient unroofing of the central Scandinavian Caledonides and record progressively younger ages of this event from SW to NE.