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The partitioning of elements amongst lower mantle phases and silicate melts is of interest in unraveling the early thermal history of the Earth. Because of the technical difficulty in carrying out such measurements, only one direct set of measurements was reported previously, and these results as well as interpretations based on them have generated controversy. Here we report what are to our knowledge only the second set of directly measured trace element partition coefficients for a natural system (KLB-1).
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The formation of Ruddlesden-Popper type layers (alternating slabs of rocksalt and perovskite structures) is seen in these oxides which is similar in many respects to what is seen in the system Sr-Ti-O. However, it was observed that there are some significant differences, for example the rocksalt and perovskite blocks in new superconducting compounds are not necessarily electrically, unlike in Sr-Ti-O systems. This will certainly render an additional coulombic bonding energy between two different types of blocks and may well lead to significant differences in their structural chemistry. In the higher order members of the various homologous series, additional Cu-O planes are inserted in the perovskite blocks. In order for the unit cell to electrically neutral the net positive charge on rocksalt block (which remains constant throughout the homologous series) should be balanced by an equal negative charge on perovskite block. It, thus becomes necessary to create oxygen vacancies in the basic perovskite structure, when width of the perovskite slab changes on addition of extra Cu-O planes. Results of atomistic simulations suggest that these missing oxygen ions allow the Cu-O planes to buckle in these compounds. This is also supported by the absence of buckling in the first member of Bi-containing compounds in which there are no missing oxygen ions and the Sr-Ti-O series of compounds. Additional results are presented on the phase stability of polytypoid structures in these crystal chemically complex systems. The studies will focus on the determination of the location of Cu(3+) in the structures of higher order members of the La-Cu-O system and whether Cu(3+) ions or oxygen vacancies are energetically more favorable charge compensating mechanism.
The aftermath of phase separation and crystal-liquid fractionation in a magma ocean should leave a planet geochemically differentiated. Subsequent convective and other mixing processes may operate over time to obscure geochemical evidence of magma ocean differentiation. On the other hand, core formation is probably the most permanent, irreversible part of planetary differentiation. Hence the geochemical traces of core separation should be the most distinct remnants left behind in the mantle and crust, In the case of the Earth, core formation apparently coincided with a magma ocean that extended to a depth of approximately 1000 km. Evidence for this is found in high pressure element partitioning behavior of Ni and Co between liquid silicate and liquid iron alloy, and with the Ni-Co ratio and the abundance of Ni and Co in the Earth's upper mantle. A terrestrial magma ocean with a depth of 1000 km will solidify from the bottom up and first crystallize in the perovskite stability field. The largest effect of perovskite fractionation on major element distribution is to decrease the Si-Mg ratio in the silicate liquid and increase the Si-Mg ratio in the crystalline cumulate. Therefore, if a magma ocean with perovskite fractionation existed, then one could expect to observe an upper mantle with a lower than chondritic Si-Mg ratio. This is indeed observed in modern upper mantle peridotites. Although more experimental work is needed to fully understand the high-pressure behavior of trace element partitioning, it is likely that Hf is more compatible than Lu in perovskite-silicate liquid pairs. Thus, perovskite fractionation produces a molten mantle with a higher than chondritic Lu-Hf ratio. Arndt and Blichert-Toft measured Hf isotope compositions of Barberton komatiites that seem to require a source region with a long-lived, high Lu-Hf ratio. It is plausible that that these Barberton komatiites were generated within the majorite stability field by remelting a perovskite-depleted part of the upper mantle transition zone.
Ca- Al-rich inclusions (CAIs) are commonly surrounded by Wark-Lovering (WL) rims - thin (approx. 50 micrometers) multilayered sequences - whose mineralogy is dominated by high temperature minerals similar to those that occur in the cores of CAIs [1]. The origins of these WL rims involved high temperature events in the early nebula such as condensation, flashheating or reaction with a nebular reservoir, or combinations of these processes. These rims formed after CAI formation but prior to accretion into their parent bodies. We have undertaken a coordinated mineralogical and isotopic study of WL rims to determine the formation conditions of the individual layers and to constrain the isotopic reservoirs they interacted with during their history. We focus here on the spinel layer, the first-formed highest- temperature layer in the WL rim sequence. Results and Discussion: We have performed mineralogical, chemical and isotopic analyses of an unusual ultrarefractory inclusion from the Allende CV3 chondrite (SHAL) consisting of an approx. 500 micrometers long single crystal of hibonite and co-existing coarsegrained perovskite. SHAL is partially surrounded by WL rim. We previously reported on the mineralogy, isotopic compositions and trace elements in SHAL [2-4]. The spinel layer in the WL rim is present only on the hibonite and terminates abruptly at the contact with the coarse perovskite. This simple observation shows that the spinel layer is not a condensate in this case (otherwise spinel would have condensed on the perovskite as well). The spinel layer appears to have formed by gas-phase corrosion of the hibonite by Mg-rich vapors such that the spinel layer grew at the expense of the hibonite. We also found that the spinel layer has the same 16Orich composition as the hibonite. The spinel layer is polycrystalline and individual crystals do not show a crystallographic relationship with the hibonite. An Al-diopside layer overlies the spinel layer, and is present on both the hibonite and perovskite. While the spinel is 16O-rich, WL-rim perovskite and pyroxene are 16O-poor. This isotopic heterogeneity likely reflects O isotopic equilibration of WL-rim perovskite and pyroxene with a planetary O isotopic reservoir after the WL rim formation. The hibonite is zoned and contains wt.% levels of Ti, Mg and Fe in contact with the Fe-bearing spinel (Sp60Hc40) in the WL rim. The Fe enrichment in spinel is likely related to the Na-Fe metasomatism that is ubiquitous in Allende. Conclusions: The petrography and microstructure of the spinel layer in a WL rim sequence shows that it formed by gas phase reactions at high temperature in the nebula. The oxygen isotopic composition of the spinel indicates that this WL rim layer formed in the same (or similar) nebular gas reservoir as the host CAI.
Introduction: Primitive meteorites contain Calcium-Aluminum-rich Inclusions (CAIs)that preserve the records of the earliest times in the Solar System. CAIs record O isotopic variations with time and/or location from solar (16O-rich) to planetary (16O-poor) compositions [e.g., 1, 2].However, CAIs were also affected by parent body processing to varying degrees. The goal of this work was to identify the processes that resulted in the fine-scale spatial variations in oxygen isotopic compositions within the most ancient Solar System solids. We report oxygen isotopic composition of CAIs from MIL 090019 CO3.1 carbonaceous chondrite. MIL 090019 contains a high abundance of CAIs showing mineralogical as well as textural variations [3]. Methods: We analyzed 10 CAIs with different mineralogical assemblages, classified as corundum-bearing, grossite-bearing, hibonite-bearing, melilite-bearing, anorthite-bearing inclusions and amoeboid olivine aggregates(AOA’s). The mineralogical and petrological characterization of CAIs was performed using the JOEL Hyperprobe 8530 electron microprobe at NASA JSC. Oxygen isotopic imaging of the CAIs was done using the Cameca NanoSIMS 50L ion microprobe at NASA JSC. We followed the analytical protocol described in [4]. O-isotopic maps of the CAI were acquired by rastering a ~3 pA primary Cs+beam at 16 keVover an area of 20 ×20μm2over a period of ~7 hours. Negative secondary ions of 16O−, 17O−, 18O−, 28Si−, 24Mg16O−, 27Al16O−, and 40Ca16O−were simultaneously acquired using electron multiplier detectors at a mass resolving power sufficient to resolve the 16OH−interference from the 17O−peak, where the contribution of 16OH−was <0.1%.An electron flood gun was used to mitigate sample charging during the analyses. We used San Carlos olivine and Madagascar hibonite as isotopic standards to correct for the instrumental mass fractionation. The O-isotopic ratios were corrected for the quasi simultaneous arrival (QSA) effect and the detector dead time. All reported errors are 1 sigma. Results: The MIL 090019 carbonaceous chondrite hosts refractory inclusions varying in their mineralogies and textures[2, 3]. We found the oxygen isotopic compositions of mineral phases in MIL 090019 CAIs show large variations with Δ17O values varying from -27 to 0‰. A corundum-bearing inclusion records isotopic variations with corundum, melilite, perovskite, and anorthite showing Δ17Ovalues of -20 ±2.5, -10.2 ± 6.5, -10.7 ± 1.3,and -1.2 ± 3.2‰ respectively. A grossite-hibonite-bearing inclusion also records O-isotopic heterogeneity, where grossite, hibonite, Mg-rich spinel, and melilite show -3.7 ± 2.7, -22.9 ± 2.5, -21.4 ± 2.6, and -17.3 ± 2.8‰ respectively, whereas perovskite grains record a heterogenous isotopic composition ranging from -22.6 to -13.9‰. A hibonite-bearing inclusion shows average Δ17Ο values of -17.6 ± 5.3 ‰ and -15.1 ± 4.8 ‰ for hibonite and spinel respectively. The two perovskite grains also show relatively 16O-poor composition (Δ17Ο = -11.8 ± 6.4 ‰ and -8.6 ± 5.2 ‰). Anorthite has ~0.1 wt% Na2O and ~0.52 wt% FeO and a Δ17Ο value of -7.2 ± 5.0 ‰, whereas diopside records a Δ17Ο value of -3.5 ± 5.0 ‰. A hibonite-spinel-rich inclusion contains spinel and hibonite that are16O-rich (Δ17Οvalues -21.3 ± 2.8‰ and -19.4 ± 3.0‰, respectively). Melilite shows heterogenous oxygen isotopic composition, with an inner 16O-rich region and an outer 16O-poor region separated by a sharp boundary, with a bulk Δ17Ο= -16.2 ± 2.1‰, whereas two perovskite grains are relatively 16O-poor (Δ17Ο= -13.2 ± 3.7‰ and -7.1 ± 3.4‰).A spinel-rich inclusion contains perovskite grains with Δ17Οranging from -27.0 to -22.8 ‰and spinel with average Δ17Οvalue of -21.4 ± 2.8 ‰. In a melilite and spinel-rich CAI, perovskite, spinel, diopside, melilite, and Al-Ti-rich pyroxene record Δ17Ο values of-26.9 ± 4.6‰, -22.4 ± 4.3‰,-23.4 ± 4.9, -19.3 ± 3.4, and -19.0 ± 4.3 ‰ respectively. Olivine and spinel components of 2 AOA’s record a 16O-rich composition with Δ17O~-23‰. Discussion: The coexistence of 16O-rich and 16O-poor minerals within CAIs has been attributed to differing degrees of O isotopic exchange, reflecting a wide range of O diffusion rates[5]. However, the O isotopic heterogeneity observed in these CAIs cannot be explained by oxygen diffusion in nebular or parent body settings alone. The isotopic imaging of CAIs from the MIL 090019 meteorite shows variations in the O isotopic composition recorded in their mineral components, suggesting that these CAIs record the oxygen isotopic heterogeneity in the nebular gas from which they condensed. The coexistence of 16O-rich and 16O-poor mineral components in the CAI-forming region suggests that the gas in this region was not well-mixed and consisted of distinct 16O-rich and 16O-poor gaseous reservoirs.