First widespread occurrence of rare phosphate chladniite in a meteorite, winonaite Graves Nunataks (GRA) 12510: Implications for phosphide–phosphate redox buffered genesis in meteorites

Electron probe microanalysis (EPMA)

Phase identification and quantitative major and minor element chemistry measurements were completed using electron probe microanalysis (EPMA) using the JEOL 8530 field emission (FE) electron microprobe at NASA Johnson Space Center (JSC) using a ZAF correction. Silicate minerals were analyzed using a 15 kV accelerating voltage, a 10 nA beam current, and a 5 μm spot size via wavelength-dispersive X-ray spectroscopy (WDS). Metals and sulfides were analyzed using a 15 kV accelerating voltage, 20 nA beam current, and a 1–5 μm spot size. Natural and synthetic standards were used, including canyon diablo troilite for S. Phosphates were analyzed using 15 kV accelerating voltage, 20 nA beam current, and 10 μm spot size. The analysis of phosphates, particularly apatite, followed procedures established and described in McCubbin et al. (2021). Natural and synthetic minerals were used as standards, including SrF₂, albite, quartz, stikin anorthite, springwater olivine, Wilberforce apatite, ilmenite, and rhodonite [crystal setup LDE1: F, TAP (Na, Si, Al, Mg), PET (P, S), PETL (Cl, Ca, Ti), and LIFH (Fe, Mn)]. Individual EPMA analyses are included in the Online Materials^[1] files.

Scanning electron microscopy (SEM)

Modal mineralogy was calculated using elemental maps produced with the JEOL 7600 scanning electron microscope (SEM) at NASA JSC. Elemental maps were collected using an accelerating voltage of 15 kV, 20 nA beam current, and 8 mm working distance at a resolution of <0.5 μm/pixel. Mineral phases were identified by the distribution of 16 elements (Na, Mg, Al, Si, P, S, K, Ca, Ti, Cr, Mn, Fe, Ni, Zn, C, O) along with energy-dispersive X-ray spectroscopy (EDS) spot analyses. Modal mineral abundances were determined by converting the elemental maps to mineralogical maps in XMapTools (Lanari et al. 2014, 2019) (Table 1; Fig. 2).

Figure 2

Mineralogical map of GRA 12510. Chladniite grains from 1–200 μm are scattered throughout the section often on margins between silicate clasts and metal. The areas immediately surrounding chladniite are mostly Mn free, but several Mn-rich chromite grains (7.89 wt% MnO) and one tiny alabandite grain (MnS) are found elsewhere in the section. Schreibersite is also anti-correlated with chladniite. Black box insets correspond to areas A and B in Figure 1. White circles are 75–150 μm LA-ICP-MS pits from previous trace element analyses.

Electron backscatter diffraction (EBSD)

Quantitative phase indexing and microstructural maps of chladniite-bearing regions were collected using the JEOL 7900F FE-SEM at NASA JSC. Electron backscatter diffraction patterns (EBSPs) were collected using an Oxford Instruments Symmetry crystal metal oxide scintillator detector and an Oxford instrument Aztec 4.1 software package. Acquisition parameters included a 20 kV accelerating voltage, 10 nA beam current, ~20 mm working distance, and a 70° tilt relative to the beam incidence. Phosphate EBSPs were indexed using a fillowite match unit after the crystal structure of Araki and Moore (1981). In addition, a merrillite match unit based on the structural data of Xie et al. (2015) was included in all maps to identify additional phosphates and ensure accurate structural indexing of the grains. Maps were collected with a step size ranging from 0.5–0.2 μm.

Post-processing of the EBSD data used Oxford Instruments Aztec Crystal 2.2 program. All EBSD data were given a wild-spike noise reduction and a seven nearest neighbor zero-solution correction. An additional correction was applied to remove systematic misindexing of the fillowite data with disorientation relation of 180° <11–20> caused by mirror symmetry normal to the axis. Forescatter electron, EBSD phase, and inverse pole figure (IPF) maps can be found in Figure 3. Fillowite pole plots are equal area, lower hemisphere projections.

Figure 3

EBSD data of three occurrences of chladniite [(a) area A, (b) area A, (c) area B] and its associated mineral assemblages with forward scattered electron diode (FSD), phase overlay, and inverse pole (IPF) images along with the Kikuchi patterns, Kikuchi bands, and Kikuchi solutions as well as pole figures {0001}, {1010}, and {1120}. (Figure continues on next page.)

Raman spectroscopy and Raman spectroscopic imaging

Chladniite and its associated minerals were measured using a WITec alph300R Raman microscope (XMB3000-3003) at NASA JSC. Both single spectra and a Raman image were collected. Incident 532 or 488 nm excitation was generated using a WITec diode laser (XSL3100-1154), producing ~170 μW or ~100 μW of incident power at the sample surface, respectively. For single spectral collections, the incident laser was focused on the sample surface using either a 100 × magnification Zeiss EC Epiplan-Neofluar objective producing a 0.9 μm beam spot on the sample surface or a 50× magnification Zeiss EC Epiplan objective producing a 1.2 μm spot size on the sample surface. The Raman scattered light was dispersed using a WITec UHTS600 (XMC3200-0600) spectrometer with a 300 g/mm grating, and the spectrum was detected using a back-illuminated, thermoelectrically cooled (–60 °C) CCD (XMC3022-1001). Single spectra were collected by averaging ~5–10 accumulations, each with a 10 s integration time. The Raman image of 100 × 100 spectra was collected over a 300 × 300 μm area of the sample for a Raman image resolution of 3 μm/pixel and a total of 10 000 individual spectra. The 50 × magnification objective was used to collect the Raman image. The accumulation time for each individual Raman spectrum in the image was 5 s for an image accumulation time of 14 h and 8 min. Principal component analysis (PCA) of the image was performed using WITec Project FIVE software to identify and map the main component minerals chladniite, merrillite, enstatite, diopside, olivine, and feldspar as shown in Figure 4a. Representative Raman spectra of chladniite and merrillite, as well as fillowite from the literature, are shown in Figure 4b.

Figure 4

(a) Raman image overlaying an optical microscope image exhibiting the distribution of chladniite and merrillite in area B of GRA 12510. (b) Raman spectra of chladniite compared with merrillite and fillowite (RRUFF ID: R110143).

Occurrence

GRA 12510 is a coarse-grained winonaite with abundant metal veins and pools separating silicate clasts that are either dominated by orthopyroxene or orthopyroxene+olivine. Modal mineralogy is 65.43% silicate (28.27% orthopyroxene, 21.63% olivine, 4.54% clinopyroxene, and 10.87% plagioclase), 30.01% metal (26.08% kamacite, 3.50% taenite, and 0.43% schreibersite), with minor troilite (2.73%), chromite (0.50%), merrillite (1.13%), chladniite (0.32%), and trace apatite (Table 1; Fig. 2). Both silicate clast lithologies have similar silicate:metal area% ratios (84.33:8.15 and 85.11:10.77 for the orthopyroxene-rich and orthopyroxene + olivine-rich clasts). However, the orthopyroxene-rich clasts have more orthopyroxene (45.05 vs. 31.68%), less clinopyroxene (5.36 vs. 11.09%), similar abundances of plagioclase (12.85 vs. 11.64%), and less olivine (21.07 vs. 30.70%). A single 5 μm diameter grain of alabandite (MnS) was also found in an orthopyroxene-rich silicate clast in contact with troilite and schreibersite. A single F-rich, Cl-bearing apatite grain was found as an inclusion in merrillite. Notably, the orthopyroxene-rich clast has substantially more phosphate (0.67% chladniite and 1.66% merrillite vs. 0.31% chladniite and 1.23% merrillite) and less phosphide (0.09% vs. 0.81% schreibersite).

Numerous 1–500 μm chladniite grains were found, often on the margins between silicate clasts and the kamacite portions of the large metal veins. The largest chladniite grains are associated with merrillite, kamacite, taenite, troilite, albite, forsterite, diopside, and enstatite (Fig. 2). Rarely, chladniite encloses <3 μm merrillite grains as seen in Figure 2b. Additionally, even though chladniite in this meteorite contains Mn, the areas immediately surrounding chladniite are mostly Mn free. However, several Mn-rich chromite grains (7.89 wt% MnO) and one tiny alabandite grain (MnS) are found elsewhere in the section.

Composition of chladniite

The average composition of 37 analyses of 7 grains of chladniite in GRA 12510 is shown in Table 2. These chladniite grains include those near metal, merrillite, and silicates, as well as inclusions in merrillite and albite. Based on these analyses, the calculated mineral formula for chladniite in GRA 12510 is Na_2.7Ca_1.25(Mgi_0.02Mn_0.69Fe_0.20)_Σ10.91(PO₄)₉ after the paper of Hatert et al. (2021), stoichiometry Na₃CaMg₁₁(PO₄)₉.

EBSD

EBSD analyses of GRA 12510 revealed that the chladniite structure agreed well with fillowite (Araki and Moore 1981), thus falling within the fillowite-type phosphate group. Chladniite areas were found using forescatter electron images (Fig. 3). Diffraction patterns were collected from one large chladniite grain in area A, as well as chladniite areas with multiple grain boundaries in area A and B (areas A and B matching those identified in BSE images in Fig. 1). The mean angular deviation values of the electron backscatter patterns for the maps ranged between 0.87 to 0.63 for fillowite-structure materials, corroborating the chladniite-type phosphate.

Crystallographic orientations displayed as IPP maps and pole plots showed a single 40 × 60 μm chladniite grain in area A (Fig. 3a). Other chemically coherent chladniite areas including a 100 × 100 μm area (Fig. 3b) and an ~500 μm diameter area (Fig. 3c) were actually revealed through EBSD to be monomineralic granular aggregates with grain sizes down to 3 μm. The pole plots do not show a systematic pattern consistent with preferred growth orientation, but rather with multiple random orientations, indicating nucleation and coalescence of grains (Prior et al. 1999).

Winonaite parent body pressure-temperature-redox conditions

There are some thermodynamic constraints on the formation conditions of the winonaite parent body that can be used to elucidate the formation of chladniite. The winonaite parent body has been estimated to be a small body with a radius of ~100 km (Benedix et al. 2005) that corresponds to a maximum pressure of 0.01 GPa (Lucas et al. 2020). The winonaite parent body is also reduced, with silicates recording oxygen fugacities (f_O2) from IW-2.3 to IW-3.2 calculated using the FeO content of olivine (quartz-iron-fayalite) and pyroxene (quartz-iron-ferrosillite) (Benedix et al. 2005). In their study, Benedix et al. (2005) note that winonaite and IAB f_O2 estimates do not fall along any traditional metal-metal oxide buffers or the temperature and pressure-dependent graphite-carbon monoxide buffer. However, this f_O2 range calculated from the silicates is consistent with the phosphide-phosphate oxidation transition of IW-3 to IW-4 designated by the schreibersite (Fe₃P)-whitlockite [Ca₃(PO₄)₂] buffer (Bindi et al. 2023; Pasek 2015) as shown in Figure 5. Recent experiments also suggest that at temperatures >900 °C, the phosphide-to-phosphate reaction occurs rapidly, with complete oxidation from phosphide within a week and a few percent reduction of phosphate on the same timescale (Feng and Pasek 2023). The co-occurrence and close association of merrillite, chladniite, apatite, and P-bearing metallic phases within GRA 12510 suggests that the f_O2 of IW-2 to IW-4 is an intrinsic property of the precursor chondritic material and the phosphate-phosphide reaction may have buffered the final winonaite and IAB iron meteorite phase assemblages.

Figure 5

Plot of oxygen fugacity (log f_O2) vs. temperature [10 000/T(K)] for IAB irons and a winonaite along with appropriate oxygen fugacity buffers adapted from Benedix et al. (2005). Oxygen fugacity determined from orthopyroxene (blue) and olivine (red) for meteorites Cad (Caddo County), CdC (Campo del Cielo), Cop (Copiapo), Lue (Lueders), Ude (Udei Station), Win (Winona) with data falling along the dashed best fit line. Winonaite and IAB iron meteorite oxygen fugacities fall along the schreibersite (Fe₃P) – whitlockite [Ca₃(PO₄)₂] buffer purple line (Pasek 2015; Bindi et al. 2023) indicating these meteorites were likely redox buffered by the present phase assemblage of phosphide (P valence state of 0) and phosphate (P valence state of 5+). Also shown are three CO-C fugacity buffers at different pressures that fall on a different slope.

The prevailing model for winonaite-IAB parent body formation is that it underwent incomplete differentiation followed by catastrophic impact, breakup, and reassembly (Benedix et al. 2000). Most recently, the proposed layered structure of the winonaite-IAB iron meteorite parent body from the surface to the core was: (1) precursor chondritic material; (2) diverse lithologies that experienced limited metamorphism and FeNi-FeS partial melting; (3) residues of silicate partial melting; and (4) incompletely differentiated molten metal represented by IAB iron meteorites (Hunt et al. 2017; Zeng et al. 2019). Adding onto this model, winonaite and IAB iron meteorites record rapid cooling rates of 0.48–1.75 °C/year, and therefore cooled from their peak magmatic temperatures as collisional fragments during breakup prior to reassembly (Anzures et al. 2022). Finally, the winonaite parent body also experienced multiple impact events that caused liberation from the parent body before impacting and being found on Earth. Cosmic-ray exposure ages for winonaites range from 0.02 to 0.08 Ga (Benedix et al. 1998), while IAB iron meteorites range from 0.4 to 1.0 Ga (Voshage 1967).

Origin of chladniite

The only other occurrences of chladniite have been as a single large grain in a complex silicate-bearing inclusion in IAB iron meteorite Carlton (McCoy et al. 1994), a minor phase in lodranite GRA 95209 (Floss 1999; Grew et al. 2010; McCoy et al. 2006; McCoy and Carlson 1998), as inclusions in merrillite-dominated phosphate aggregate in a IIE iron meteorite Elga (Litasov and Podgornykh 2017), in a couple of terrestrial pegmatites in Argentina (Vallcorba et al. 2017) and Brazil (Hatert et al. 2021), and as a REE-bearing Y-chladniite in a granulite from Antarctica (Grew et al. 2006). The first terrestrial identification of chladniite was found in a pegmatite from Córdoba, Argentina, as up to 200 μm diameter inclusions in beusite, the Mn analog of the graftonite phosphate structural group (Vallcorba et al. 2017). Based on chemical and textural evidence, the association of beusite and chladniite was interpreted as a replacement product of garnet during the magmatic stage (Vallcorba et al. 2017) caused by destabilization of garnet due to phosphorous buildup (Colombo et al. 2012). Chladniite has also been found in another pegmatite in Sapucaia, Brazil (Hatert et al. 2021). Y-chladniite was also found in a granulite in Larsemann Hills, Antarctica as up to 250–1000 μm inclusions in fluorapatite associated with wagnerite, xenotime-(Y), monazite-(Ce), P-bearing K-feldspar, biotite, sillimanite, quartz, and pyrite. Y-chladniite was inferred to have formed at 800–860 °C and 0.6–0.7 GPa by reaction of biotite with an anatectic melt with phosphorous buildup by interaction with fluorapatite (Grew et al. 2006). However, terrestrial chladniite likely has a different mode of formation than meteoritic chladniite because of the much higher pressures and more oxidizing environments on Earth, as well as the inferred differences in geological processes that occur on small asteroids vs. the Earth’s crust.

For meteoritic chladniite formation, two hypotheses for its formation are either as a: (1) replacement product of a phosphate [either graftonite (Floss 1999) or panethite (Litasov and Podgornykh 2017)], which originated by reaction of olivine and orthopyroxene with metallic iron containing P or (2) crystallization from the reaction of plagioclase and pyroxene with metallic iron or schreibersite (Floss 1999). In lodranite GRA 95209, chladniite is also associated with a graftonite structural group phosphate in the iron analog graftonite as <50 μm thick veins intersecting silicate phases adjacent to remnant metal or troilite (Floss 1999). However, in IAB iron meteorite Carlton, the single chladniite grain (175 × 975 μm) is not associated with a graftonite group phosphate, but rather with chlorapatite. In IIE iron meteorite Elga, chladniite was found in only one site as <10 μm inclusions in merrillite surrounded by schreibersite rims and enclosed within kamacite-taenite metal (Litasov and Podgornykh 2017). In Elga, chladniite was suggested to be poorly crystalline due to its Raman spectra always being contaminated by merrillite peaks (Litasov and Podgornykh 2017).

Chladniite in winonaite GRA 12510 is much more abundant and found in larger grain sizes here compared with other meteorites. Numerous 1–500 μm chladniite grains were found, often on the margins between silicate clasts and the kamacite portions of the large metal veins. The largest chladniite grains are associated with merrillite, kamacite, taenite, troilite, albite, forsterite, diopside, and enstatite (Fig. 2). Additionally, unlike Elga where a few tiny poorly crystallized chladniite grains are included in merrillite (Litasov and Podgornykh 2017), sometimes chladniite in GRA 12510 encloses <3 μm merrillite grains as seen in Figure 2b.

Chladniite compositions differ widely, which suggests that the composition is controlled by the geologic environment of formation rather than by crystal-chemical constraints (Vallcorba et al. 2017). Terrestrial pegmatitic chladniite has high Mn and Fe (13.96–14.42 wt% MnO and 15.98–17.37 wt% FeO) reflective of its pegmatitic environment (Hatert et al. 2021; Vallcorba et al. 2017) and association with Mn-phosphate beusite in Argentina (Vallcorba et al. 2017). Terrestrial granulite chladniite has similar Fe content (15.88 wt% FeO) but virtually no Mn (0.26 wt% MnO) due to its almost Mn-free environment (Grew et al. 2006). The chladniite of lodranite GRA 95209 is also rich in Fe with a little less Mn [8.61 wt% MnO and 14.4 wt% FeO (Floss 1999)] reflective of its association with Fe-phosphate graftonite. Chladniite in silicate clasts of iron meteorites has very low Mn and Fe, including IAB iron Carlton (0.30 MnO and 2.23 wt% FeO, respectively) (McCoy et al. 1994) and IIE Elga (1.37 MnO and 5.11 wt% FeO) (Litasov and Podgornykh 2017). GRA 12510s chladniite also has relatively low overall Mn and Fe (3.92 wt% MnO and 1.14 wt% FeO, respectively). However, GRA 12510s chladniite is 10 times higher in Mn than in Carlton, which could have been scavenged from MnS or Mn-rich chromite that are spatially anti-correlated with abundant chladniite in the sample. Interestingly, this composition plots slightly within the proposed miscibility gap in a Fe²⁺-Mn-Mg ternary for fillowite-type phosphates (Fransolet et al. 1998; Hatert et al. 2021), as shown in Figure 6. This possible miscibility gap has been hypothesized because of the larger difference in ionic radii between Mg and Mn compared with their similarities to Fe²⁺, although this could also be due to small sample bias with only seven total occurrences of chladniite.

Figure 6

Ternary Fe²⁺-Mn-Mg diagram of natural fillowite-type phosphates. The most Mn-rich analysis of chladniite in GRA 12510 falls slightly within the proposed miscibility gap of Hatert et al. (2021).

Altogether, chladniite appears to form alongside other phosphates, with their chemistries reflecting the diverse environment of their formation. Meteoritic chladniite likely formed through subsolidus oxidation of schreibersite (Fe,Ni)₃P scavenging Na from albite, Ca from diopside, Mg from enstatite/forsterite, Fe from kamacite/taenite, and Mn from alabandite/chromite when available. Near Mg-end-member enstatite and forsterite (Mg# >90) are generally the most abundant mineral aside from metal in winonaites, IAB irons, and lodranites, consistent with the Mg-, rather than Fe- or Mn-, end-member of the fillowite group being present in GRA 12510. Whereas Carlton likely formed in a Cl-rich environment with chlorapatite present (McCoy et al. 1994), and GRA 95209 formed in a Cl-poor and Fe-rich environment (Floss 1999). Although Elga has both chladniite, merrillite, and schreibersite like GRA 12510, its silicates appear to be more oxidized with an average orthopyroxene composition of En₇₆Fs_21.5Wo_2.5. Elga’s lower Mn in chladniite compared with GRA 12510 is also reflected in its Mn-poor chromite (2.15 vs. 7.89 wt% MnO) and the presence of only troilite rather than troilite and alabandite, respectively.