Band 106, Heft 1

Pressure-volume-temperature ( P-V-T ) data of synthetic Mg 7 Si 2 O 8 (OH) 6 phase A were collected under P-T conditions up to ~10.5 GPa and 900 K by energy-dispersive X‑ray diffraction using a cubic type multi-anvil apparatus, MAX80, located at the Photon Factory–Advanced Ring (PF-AR) at the High Energy Accelerator Research Organization (KEK). P-V EoS using only room-temperature data yielded V 0 = 511.6(2) Å 3 , K T0 = 106.8(18) GPa, and pressure derivative KT′ =$K_{T}^{\prime}=$ 3.88(38). These parameters were consistent with the subsequent equation of state (EoS) analysis. The compressibility of phase A was anisotropic, with its a -axis being ~26% more compressible than the c -axis, which is normal to the plane of the distorted close-packed layers. A fit of the present data to the high-temperature Birch-Murnaghan EoS yielded V 0 = 511.7(3) Å 3 , K 0 = 104.4(24) GPa, K ′ = 4.39(48), (∂ K T /∂ T ) P = –0.027(5) GPa K –1 , and thermal expansion α = a + b T with values of a = 2.88(27) × 10 –5 K –1 and b = 3.54(68) × 10 –8 K –2 . The lattice dynamical approach by the Mie-Grüneisen-Debye EoS yielded θ 0 = 928(114) K, q = 2.9(10), and γ 0 = 1.19(8). The isobaric heat capacity C P of phase A at 1 atm. was calculated based on the Mie-Grüneisen-Debye EoS fit of present P-V-T data. In addition, the density profiles of subducting slabs with different degrees of serpentinization were also calculated along the cold geotherm up to ~13 GPa. The serpentinization of subducting slab will significantly lower the density of slab at shallower depth; however, this effect becomes negligible when antigorite dehydrates to phase A. Because the phase A bearing subducting slab is supposed to be denser than the surrounding mantle, the water can transport into deeper parts of the upper mantle and the transition zone.

X‑ray diffraction analysis and ultrasonic measurements of a glass with the pyrope composition were conducted to determine its structural and elastic properties at pressures from 1 atm to 12.9 GPa. Our results indicate that its structural evolution is closely related to changes in the compression wave velocity ( V P ), shear wave velocity ( V S ), and Poisson ratio. We observed three modes of pyrope glass compression. Moderate shrinkage in the intermediate-range ordered structure occurred at pressures below 6 GPa. Significant shrinkage in the intermediate-range ordering was observed at pressures between 6 and 9 GPa. We observed changes in the short-range ordered structure at pressures above 9 GPa, which were associated with an increase in the coordination number of tetrahedral cations. The absolute values of V P and V S in pyrope glass are similar to those in magnesium-bearing silicate glasses with enstatite and diopside compositions. However, the velocities are higher than those observed in sodium aluminum silicate glasses with jadeite and albite compositions. This indicates that the velocities are governed by the initial density of a glass, which is determined by its chemical composition. In terms of pressure, the velocity minimum in pyrope glass occurs at ~5 GPa, which is similar to the velocity minima in fully polymerized glasses, such as jadeite and albite. The degree of polymerization in pyrope glass is intermediate, and it has a relatively polymerized network. A drastic increase in velocity was observed when the pyrope glass was subjected to pressures above 7–8 GPa, and the velocity exceeded that observed in silicate glasses. Densification phenomenon, such as an increase in the Al coordination number, was efficiently promoted. This was because the cationic field strength of Mg 2+ exceeds those of typical non-network forming cations. Magnesium cations may have an important role in controlling the behavior of silicate glass, and partially melted mantle becomes enriched with Mg under pressure. Studying Mg-bearing aluminosilicate glasses can thus help us to better understand the behavior of magma deep in the interior of the Earth.

Birnessite-like minerals are among the most common Mn oxides in surficial soils and sediments, and they mediate important environmental processes (e.g., biogeochemical cycles, heavy metal confinement) and have novel technological applications (e.g., water oxidation catalysis). Ca is the dominant interlayer cation in both biotic and abiotic birnessites, especially when they form in association with carbonates. The current study investigated the structures of a series of synthetic Ca-birnessite analogs prepared by cation-exchange with synthetic Na-birnessite at pH values from 2 to 7.5. The resulting Ca-exchanged birnessite phases were characterized using powder X‑ray diffraction and Rietveld refinement, Fourier transform infrared spectroscopy, Raman spectroscopy, X‑ray photoelectron spectroscopy, and scanning and transmission electron microscopy. All samples synthesized at pH values greater than 3 exhibited a similar triclinic structure with nearly identical unit-cell parameters. The samples exchanged at pH 2 and 3 yielded hexagonal structures, or mixtures of hexagonal and triclinic phases. Rietveld structure refinement and X‑ray photoelectron spectroscopy showed that exchange of Na by Ca triggered reduction of some Mn 3+ , generating interlayer Mn 2+ and vacancies in the octahedral layers. The triclinic and hexagonal Ca-birnessite structures described in this study were distinct from Na- and H-birnessite, respectively. Therefore, modeling X‑ray absorption spectra of natural Ca-rich birnessites through mixing of Na- and H-birnessite end-members will not yield an accurate representation of the true structure.

A variety of experimental techniques have been proposed to measure the composition of aqueous fluids in high-pressure experiments. In particular, the “diamond trap method,” where the fluid is sampled in the pore space of diamond powder and analyzed by laser-ablation ICP-MS after the experiment, has become a popular tool. Here, we carried out several tests to assess the reliability of this method. (1) We prepared several capsules loaded with fluid of known composition and analyzed the fluid by laser-ablation ICP-MS, either (a) after drying the diamond trap at ambient condition; (b) after freezing and subsequent freeze-drying; and (c) after freezing and by analyzing a frozen state. Of these methods, the analysis in the frozen state (c) was most accurate, while the results from the other two methods were poorly reproducible, and the averages sometimes deviated from the expected composition by more than a factor of 2. (2) We tested the reliability of the diamond trap method by using it to measure mineral solubilities in some well-studied systems at high pressure and high temperature in piston-cylinder runs. In the systems quartz-H 2 O, forsterite-enstatite-H 2 O, and albite-H 2 O, the results from analyzing the diamond trap in a frozen state by laser-ablation ICP-MS generally agreed well with the expected compositions according to literature data. However, in the systems corundum-H 2 O and rutile-H 2 O, the data from the analysis of the diamond trap were poorly reproducible and appeared to indicate much higher solubilities than expected. We attribute this not to some unreliability of the analytical method, but instead to the fact that in these systems, minor temperature gradients along the capsule may induce the dissolution and re-precipitation of material during the run, which causes a contamination of the diamond trap by solid phases. (3) We carried out several tests on the reliability of the diamond trap to measure fluid compositions and trace element partition coefficients in the eclogite-fluid system at 4 GPa and 800 °C using piston-cylinder experiments. The good agreement between “forward” and “reversed” experiments—with trace elements initially either doped in the solid starting material or the fluid—as well as the independence of partition coefficients on bulk concentrations suggests that the data obtained are reliable in most cases. We also show that the rate of quenching/cooling has little effect on the analytical results, that temperature oscillations during the run can be used to enhance grain growth, and that well-equilibrated samples can be obtained in conventional piston-cylinder runs. Overall, our results suggest that the diamond trap method combined with laser-ablation ICP-MS in frozen state yields reliable results accurate within a factor of two in most cases; however, the precipitation of accessory minerals in the diamond trap during the run may severely affect the data in some systems and may lead to a gross overestimation of fluid concentrations.

Earth’s lower mantle most likely mainly consists of ferropericlase, bridgmanite, and a CaSiO 3 - phase in the perovskite structure. If separately trapped in diamonds, these phases can be transported to Earth’s surface without reacting with the surrounding mantle. Although all inclusions will remain chemically pristine, only ferropericlase will stay in its original crystal structure, whereas in almost all cases bridgmanite and CaSiO 3 -perovskite will transform to their lower-pressure polymorphs. In the case of perovskite structured CaSiO 3 , the new structure that is formed is closely related to that of walstromite. This mineral is now approved by the IMA commission on new minerals and named breyite. The crystal structure is triclinic (space group: P 1 ) with lattice parameters a 0 = 6.6970(4) Å, b 0 = 9.2986(7) Å, c 0 = 6.6501(4) Å, α = 83.458(6)°, β = 76.226(6)°, γ = 69.581(7)°, and V = 376.72(4) Å. The major element composition found for the studied breyite is Ca 3.01(2) Si 2.98(2) O 9 . Breyite is the second most abundant mineral inclusion after ferropericlase in diamonds of super-deep origin. The occurrence of breyite has been widely presumed to be a strong indication of lower mantle (>670 km depth) or at least lower transition zone (>520 km depth) origin of both the host diamond and the inclusion suite. In this work, we demonstrate through different formation scenarios that the finding of breyite alone in a diamond is not a reliable indicator of the formation depth in the transition zone or in the lower mantle and that accompanying paragenetic phases such as ferropericlase together with MgSiO 3 are needed.

Taking into account recent publications, we provide additional comprehensive evidence that type Ib cuboctahedral diamonds and some other microcrystalline diamonds from Kamchatka volcanic rocks and alluvial placers cannot be natural and undoubtedly represent synthetic materials, which appear in the natural rocks by anthropogenic contamination. The major arguments provided in favor of the natural origin of those diamonds can be easily disproved. They include the coexistence of diamond and deltalumite from Koryaksky volcano; coexistence with super-reduced corundum and moissanite, Mn-Ni silicide inclusions, F-Cl enrichment and F/Cl ratios, and carbon and nitrogen isotopes in Tolbachik diamonds, as well as microtwinning, Mn-Ni silicides, and other inclusions in microcrystalline diamond aggregates from other Kamchatka placers. We emphasize the importance of careful comparison of unusual minerals found in nature, which include type Ib cuboctahedral diamonds and super-reduced phase assemblages resembling industrial slags, with synthetic analogs. The cavitation model proposed for the origin of Tolbachik diamonds is also unreliable since cavitation has only been shown to cause the formation of nanosized diamonds only.

Diamonds have long been mined from alluvial terrace deposits within the rainforest of Guyana, South America. No primary kimberlite deposits have been discovered in Guyana, nor have there been previous studies on the mineralogy and origin of the diamonds. Paleoproterozoic terranes in Guyana are prospective to diamond occurrences because the most productive deposits are associated spatially with the eastern escarpment of the Paleoproterozoic Roraima Supergroup. Geographic proximity suggests that the diamonds are detrital grains eroding from the <1.98 Ga conglomerates, metamorphosed to zeolite and greenschist facies. The provenance and paragenesis of the alluvial diamonds are described using a suite of placer diamonds from different locations across the Guiana Shield. Guyanese diamonds are typically small, and those in our collection range from 0.3 to 2.7 mm in diameter; octahedral and dodecahedral, with lesser cubic and minor macle forms. The diamonds are further subdivided into those with abraded and non-abraded surfaces. Abraded diamonds show various colors in cathodoluminescence, whereas most non-abraded diamonds appear blue. In all populations, diamonds are predominantly colorless, with lesser brown to yellow and very rare white. Diamonds are predominantly Type IaAB and preserve moderate nitrogen aggregation and total nitrogen concentrations ranging from trace to ~1971 ppm. The kinetics of nitrogen aggregation indicate mantle-derived residence temperatures of 1124 ± 100 °C, assuming residence times of 1.3 and 2.6 Ga for abraded and non-abraded diamonds, respectively. The diamonds are largely sourced from the peridotitic to eclogitic lithospheric upper mantle based on both δ 13 C values of –5.82 ± 2.45‰ (VPDB-LSVEC) and inclusion suites predominantly comprised of forsterite, enstatite, Cr-pyrope, chromite, rutile, clinopyroxene, coesite, and almandine garnet. Detrital, accessory minerals are non-kimberlitic. Detrital zircon geochronology indicates diamondiferous deposits are predominantly sourced from Paleoproterozoic rocks of 2079 ± 88 Ma.

A correlation between methanogenesis and dolomite formation has been reported; however, the mechanism underlying this association is not fully understood. In this study, we conducted forced carbonate precipitation experiments at room temperature in calcite-seeded Ca/Mg carbonate solutions containing either purified non-living biomass or bound extracellular polymeric substances (EPS) of the methanogen Methanosarcina barkeri . Purified non-living biomass and bound EPS was used so as to avoid the possible influence of the complex components of the growing microbial culture on carbonate crystallization. Our results demonstrated that non-living biomass of M. Barkeri can enhance the Mg incorporation into calcitic structure and induce the crystallization of disordered dolomite. In the presence of ~113 mg L –1 of non-living biomass, disordered dolomite with ~41 and 45 mol% of MgCO 3 was precipitated in solutions with initial Mg:Ca ratios of 5:1 and 8:1, respectively. A systematic increase in the MgCO 3 contents of the precipitated Ca-Mg carbonates was also observed with the increased non-living biomass concentration. Bound EPS was shown to be the component of non-living biomass that catalyzed the precipitation of disordered dolomite. At only ~25 mg L –1 of bound EPS, disordered dolomite with ~47 and 48 mol% of MgCO 3 was precipitated in solutions with initial Mg:Ca ratios of 5:1 and 8:1, respectively. We propose that adsorption of bound EPS to growing carbonate surfaces through hydrogen bonding is the key to catalyzing disordered dolomite crystallization, and that this mechanism is also applicable to natural EPS-induced dolomite formation. This study provides significant insight into the formation mechanism of microbial-induced dolomite with high δ 13 C values.

Pyrrhotites, characterized by the chemical formula Fe 1–δ S (0 < δ ≤ 1/8), represent an extended group of minerals that are derived from the NiAs-type FeS aristotype. They contain layered arrangements of ordered Fe vacancies, which are at the origin of the various magnetic signals registered from certain natural rocks and can act as efficient electrocatalysts in oxygen evolution reactions in ultrathin form. Despite extensive studies over the past century, the local structural details of pyrrhotite superstructures formed by different arrangements of Fe vacancies remain unclear, in particular at the atomic scale. Here, atomic-resolution high-angle annular dark-field imaging and nanobeam electron diffraction in the scanning transmission electron microscope are used to study natural pyrrhotite samples that contain commensurate 4C and incommensurate 4.91 ± 0.02C constituents. Local measurements of both the intensities and the picometer-scale shifts of individual Fe atomic columns are shown to be consistent with a model for the structure of 4C pyrrhotite, which was derived using X-ray diffraction by Tokonami et al. (1972). In 4.91 ± 0.02C pyrrhotite, 5C-like unequally sized nano-regions are found to join at anti-phase-like boundaries, leading to the incommensurability observed in the present pyrrhotite sample. This conclusion is supported by computer simulations. The local magnetic properties of each phase are inferred from the measurements. A discussion of perspectives for the quantitative counting of Fe vacancies at the atomic scale is presented.

Recent studies have identified gold nanoparticles in ores in a range of deposit types, but little is known about their formation processes. In this contribution, gold-bearing magnetite from the well-documented, world-class Beiya Au deposit, China, was investigated in terms of microstructure and crystallography at the nanoscale. We present the first three-dimensional (3D) focused ion beam/scanning electron microscopy (FIB/SEM) tomography of the distribution of gold nanoparticles in nanopores in the low-Si magnetite. The porous low-Si magnetite, which overprints an earlier generation of silician magnetite, was formed by a coupled dissolution-reprecipitation reaction (CDRR). The extrinsic changes in thermodynamic conditions (e.g., S content and temperature) of the hydrothermal fluids resulted in the CDRR in magnetite and the disequilibrium of Au-Bi melts. The gold nanoparticles crystallized from Au-supersaturated fluids originating from the disequilibrium of Au-Bi melts and grew in two ways depending on the intrinsic crystal structure and pore textures: (1) heteroepitaxial growth utilizing the (111) lattice planes of magnetite, and (2) randomly oriented nucleation and growth. Therefore, this study unravels how intrinsic and extrinsic factors drove the formation of gold nanoparticles at fluid-mineral interfaces.

Seaborgite (IMA2019-087), LiNa 6 K 2 (UO 2 )(SO 4 ) 5 (SO 3 OH)(H 2 O), is a new mineral species from the Blue Lizard mine, Red Canyon, San Juan County, Utah, U.S.A. It is a secondary phase found on gypsum in association with copiapite, ferrinatrite, ivsite, metavoltine, and römerite. Seaborgite occurs in sprays of light-yellow, long flattened prisms or blades, up to about 0.2 mm in length. Crystals are elongated on [100], flattened on {010}, and exhibit the forms {100}, {010}, {001}, and {101}. The mineral is transparent with vitreous luster and very pale-yellow streak. It exhibits bright lime-green fluorescence under a 405 nm laser. The Mohs hardness is ~2½. The mineral has brittle tenacity, curved or conchoidal fracture, and one good cleavage on {100}. The measured density is 2.97(2) g/cm 3 . The mineral is immediately soluble in H 2 O at room temperature. The mineral is optically biaxial (–), α = 1.505(2), β = 1.522(2), γ = 1.536(2) (white light); 2 V meas = 85(1)°; moderate r < v dispersion; orientation X ^ a ≈ 10°; pleochroic X colorless, Y and Z light green-yellow; X < Y ≈ Z . EPMA and LA-ICP-MS analyses of seaborgite undermeasured its Li, K, and Na. The empirical formula using Li, Na, and K based on the structure refinement is Li 1.00 Na 5.81 K 2.19 (UO 2 )(SO 4 ) 5 (SO 3 OH)(H 2 O). Seaborgite is triclinic, P 1, a = 5.4511(4), b = 14.4870(12), c = 15.8735(15) Å, α = 76.295(5), β = 81.439(6), γ = 85.511(6)°, V = 1203.07(18) Å 3 , and Z = 2. The structure ( R 1 = 0.0377 for 1935 I > 2σ I ) contains [(UO 2 ) 2 (SO 4 ) 8 ] 4– uranyl-sulfate clusters that are linked into a band by bridging LiO 4 tetrahedra. The bands are linked through peripheral SO 4 tetrahedra forming a thick heteropolyhedral layer. Channels within the layers contain a K site, while an additional K site, six Na sites, and an SO 3 OH group occupy the space between the heteropolyhedral layers.

Understanding the formation of high-silica rhyolites (HSRs, SiO 2 > 75 wt%) is critical to revealing the evolution of felsic magma systems and magma chamber processes. This paper addresses HSR petrogenesis by investigating an integrated data set of whole-rock geochemistry, geochronology, and mineral composition of the ~74 Ma Nuocang HSR (SiO 2 = 74.5–79.3 wt%) from the Coqen region in southern Tibet. Cathodoluminescence (CL) images show that zircons from the Nuocang HSRs can be divided into two textural types: (1) those with dark-CL cores displaying resorption features and overgrown by light-CL rims, and (2) those comprising a single light-CL zone, without dark-CL cores. In situ single-spot data and scanning images demonstrate that these two types of zircon have similar U-Pb ages (~74 Ma) and Hf isotopic compositions [ε Hf (t) = –9.09 to –5.39], indicating they were generated by the same magmatic system. However, they have different abundances of trace elements and trace element ratios. The dark-CL cores are likely crystallized from a highly evolved magma as indicated by their higher U, Th, Hf, Y, and heavy rare earth elements concentrations, lower Sm/Yb ratio, and more negative Eu anomalies. In contrast, the uniformly light-CL zircons and the light-CL rims are likely crystallized from less evolved and hotter magma, as indicated by their lower U-Th-REE abundances and higher Ti-in-zircon temperatures. This is consistent with the Ti-in-quartz geother-mometer in quartz phenocrysts that reveals that the light-CL zones are hotter than dark-CL cores. We propose that the composition and temperature differences between cores and rims of zircons and quartz record a recharge and reheating event during the formation of the Nuocang HSRs. This implies that HSR is a result of mixing between a hotter, less evolved silicic magma and a cooler, highly evolved, and crystal-rich mush. This study shows that zircon and quartz with distinct internal textures can be combined to disentangle the multi-stage evolution of magma reservoirs, providing critical insights into the origin of HSRs.

The crystal chemistry and thermal behavior of Fe-carpholite from the Pollino Massif have been investigated by a multi-method approach. A combination of optical microscopy, scanning electron microscopy, mRaman spectroscopy, thermal analysis, room-temperature single-crystal X‑ray diffraction, and high-temperature X‑ray powder diffraction was employed. Field and micromorphological observations showed that the studied carpholite occurs in veins embedded in fine-grained matapelites and coexist with quartz, calcite, chlorite, and phengite. In particular, the tiny carpholite crystals are closely associated with quartz, suggesting simultaneous formation. Structure refinements from single-crystal X‑ray diffraction confirm that carpholite crystallizes in the Ccce space group. Anisotropic refinements converged at 2.3 ≤ R (%) ≤ 2.6 and yielded unit-cell parameters a ~13.77 Å, b ~20.16 Å, c ~5.11 Å, and V ~1419 Å 3 . An X Fe [i.e., the molar fraction Fe 2+ /(Mg+Fe 2+ +Mn)] of ~0.6 was derived from the refined occupancy at the M 1 site and is correlated to structural expansion mainly along the b and a axes and to geometrical distortions of the M 1, M 2, and M 3 octahedra. mRaman spectrum of unoriented Fe-carpholite crystals exhibits several bands in the 200–1200 cm –1 region, a strong peak at 3630 cm –1 and a weak peak at 3593 cm –1 , the latter two of which account for the presence of two independent OH groups, as also revealed by the X‑ray structure refinement. The TG curve indicates a total mass loss of 15.6% in the temperature range 30–1000 °C, and the DTA curve shows a broad endothermic band at ~400 °C, extending up to ~650 °C, and weak exothermic peaks at ~700 and 750 °C. The latter may be ascribed to the breakdown of the Fe-carpholite structure and crystallization of new phases. The in situ high-temperature X‑ray powder diffraction from 30 to 1105 °C revealed no significant changes in XRD patterns from 30 to 355 °C but reflection splittings from 380 °C due to a Fe-oxidation/deprotonation process. The carpholite and deprotonated carpholite phases coexist in the temperature range 380–580 °C, whereas only the deprotonated phase is observed up to 630 °C. Above this temperature, the carpholite structure collapses and the characteristic peaks of spinel and quartz phases are observed. At 1105 °C, spinel, mullite, garnet, cristobalite, and tridymite can be clearly identified. Our results provide insight into the thermal stability of Fe-carpholites and may help understand the thermal evolution of HP/LT metasediments.

Mineralogical distribution, textures, electron probe microanalysis of visible gold, laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS) trace element analysis of pyrite, and LA-multicollector (MC-)ICP-MS sulfur isotope analysis of sulfide minerals are examined in an ore zone extending obliquely to –4 km depth in the Sanshandao gold deposit Jiaodong, China. We relate these results to the temporal and spatial ore-forming processes in the deposit to further elucidate the controls on the deposition of visible Au and fineness variation. Two generations of Au mineralization are identified. The early generation is represented by beresitization and quartz-pyrite veins in which visible Au grains are associated with pyrite (Py1 and Py2) and are characterized by high fineness [729–961; fineness = 1000×Au/(Au+Ag)]. Py1 and Py2 are both enriched in Co, Ni, and Bi and depleted in As and Au. Texturally, gold and pyrite are pristine crystals, homogeneous in composition. These features are attributed to the sulfidation of the granitic wallrock (fluid/rock interaction) that effectively destabilizes Au in the ore-forming fluids during pyrite deposition. Fineness decreases continuously from 870 at –2650 m depth to 752 at –420 m depth. The Co and Ni contents of Py1 and Py2 decrease significantly from –4000 m to –420 m depth, whereas the As contents increase. The mean δ 34 S values of Py1 increase from 10.5 to 11.8‰. The spatial variations are interpreted to be related to gradual cooling, decompression, and an enhanced degree of fluid/rock interaction with decreasing depth, which facilitated the initiation of visible gold mineralization at ca. –2700 m depth. The late generation of Au mineralization is represented by quartz-polysulfide veins in which visible Au grains are associated with multiple sulfide minerals (Py3, galena, chalcopyrite, arsenopyrite, and sphalerite). It is characterized by low fineness (549–719), and heterogeneous textures with Ag-rich parts (218–421). Py3, occurring as the rim of pyrite grain, is interpreted to form by replacement via a dissolution-reprecipitation reaction. Py3 is distinctly enriched in As (median of 10 000 ppm) and Au (2.2 ppm), but depleted in Co, Ni, and Bi. The δ 34 S values of the polysulfide minerals decrease sharply by 4 to 5‰ at depths from –1909 to –1450 m. These features are interpreted to be generated by significant decompression and phase separation of fluid, where most ore elements (e.g., Au, Ag, As, and base metal elements) are destabilized. Our study suggests that remobilization did not affect the generation of visible Au mineralization at Sanshandao.

The classification and nomenclature of mineral species is regulated by the Commission on New Minerals, Nomenclature and Classification of the International Mineralogical Association (IMA-CNMNC). This mineral species classification is necessary for Earth Sciences, as minerals constitute most planetary and interstellar materials. Hazen (2019) has proposed a classification of minerals and other Earth and planetary materials according to “natural clustering.” Although this classification is complementary to the IMA-CNMNC mineral classification and is described as such, there are some unjustified criticisms and factual errors in the comparison of the two schemes. It is the intent of the present comment to (1) clarify the use of classification schemes for Earth and planetary materials, and (2) counter erroneous criticisms or statements about the current IMA-CNMNC system of approving proposals for new mineral species and classifications.

I welcome the “Comment” from Hatert et al. (2021) related to the proposal for an “Evolutionary system of mineralogy” (Hazen 2019) and thank them for their historically informed, conceptually nuanced, and consistently constructive contribution. They offer corrections related to two facets of my paper that seemed unfairly to criticize aspects of the International Mineralogical Association’s Commission on New Minerals, Nomenclature and Classification (IMA-CNMNC) protocols for classifying minerals. First, they note an unfortunate inferred ambivalence with respect to the relationship between the IMA system and the new evolutionary system. If I was once ambivalent, that view has changed. Having spent the past two years in an ongoing effort to develop this new historical approach, I am struck every day by the power of the IMA-CNMNC system of species classification and nomenclature, which is fundamental and indispensable to the science of mineralogy . As Hatert et al. suggest, any new approach to organizing natural solids, including one focused on planetary evolution, must rest on the foundation provided by the IMA-CNMNC and its many volunteers who selflessly bring order to the mineral kingdom. In the best scenario, the evolutionary system may one day emerge as one of several useful approaches that complement and amplify but in no way replace this core IMA-CNMNC foundation, as clearly stated in the abstract of Hazen (2019). Second, Hatert et al. (2021) offer corrections regarding the IMA-CNMNC approach to classification, in particular a mischaracterization of the formal process to incorporate amorphous phases, poorly crystalline materials, and loosely defined “mineraloids.” I am grateful for the clarifications, as well as the implication that IMA protocols may facilitate the embrace of additional such phases in the future. Finally, I welcome the chance to explore further the emerging concept of “natural kinds” as applied to the mineral kingdom. Here, our thoughts differ. I suggest that minerals, considered in their information-rich, idiosyncratic, paragenetic contexts (in contrast to IMA-CNMNC species), have the potential to represent quintessential examples of “natural kinds.” Furthermore, when viewed in their evolutionary context, minerals offer an intriguing opportunity to expand the concept of “historical natural kinds” beyond its present limited and, at times, controversial use in biology, into the realm of the co-evolving geosphere and biosphere.

In this issue This New Mineral Names has entries for 14 new species, including amamoorite, ammoniomathesiusite, aravaite, arsenowagnerite, fluorarrojadite-(BaNa), fluorcarmoite-(BaNa), folvikite, gadolinite-(Nd), goryainovite, kruijenite, natrowalentaite, nollmotzite, qatranaite, and triazolite.