Pallasite, MG (main-group)
high-Δ17O subgroup
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Found 1820
24° 12.2′ S., 68° 48.4′ W. In 1828, some small Imilac specimens were obtained on behalf of the British and Royal Scottish Museums in Buenos Aires from an Indian, José Maria Chaile. He had found the first specimens in the Atacama Desert southwest of Imilac, Chile in about 1820, and had traveled through the Atacama Desert and the Andes Mountains to sell the specimens in the capital city. In January of 1854, a professor in Santiago named Philippi was shown the strewnfield by Chaile, where he recovered numerous small specimens weighing ~4.5 kg. He also identified a hole 6 m deep thought to have been excavated by Indians searching for the supposed metallic vein. The largest mass of 198.1 kg was purchased from an Indian for the British Museum in 1877. In the intervening years thousands of smaller fragments were recovered such as the intricately patterned specimen pictured below weighing only 4.0 g.

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These pallasite masses have been perfectly preserved in the extremely dry environment of the Andes Mountains. The meteorite is composed of equal parts olivine and FeNi-metal. The yellow- to orange-colored, angular, highly fractured olivine crystals have an average size of 10 mm, but some are twice as large. The metal in the smaller specimens shows evidence of violent shearing and deformation, with frictional heating reaching recrystallization temperatures. No Thomson (Widmanstätten) structure is present on etched sections.

Recent work by Killgore and McHone (1997), using modern navigation aids, has revealed the existence of a pattern of rays of fragments emanating from the east side of an 8-m impact pit. Two smaller depressions located in line with the large pit show evidence which suggests they also were formed by an impact from an object approaching from the west-southwest, defining a strewnfield of 400 m × 200 m. Erosional forces transported many of the smaller masses downhill to the southeast.

Three tenable scenarios for the formation of the main-group pallasites are presented here, while other plausible hypotheses are outlined below. The first scenario utilizes a passive mechanism to explain the silicate–metal mixing, the second envisions an impact-induced injection of molten metal into olivine at a near-surface location, and the third proposes that a glancing impact disrupted a smaller body, which was followed by its reassembly into a pallasite parent body. Scenario 1 (e.g., Boesenberg et al., 2012; Donohue et al., 2018)

  1. Olivine crystallized from the silicate liquid at the lowest layer of the mantle, the core–mantle interface.
  2. Cooling and contraction of the metallic core produced a 2% volume reduction leaving a void at the core–mantle boundary.
  3. The overlying crystalline olivine then collapsed into the viscous metal where heating and mixing occurred to produce the pallasitic structure.
  4. Rounding of olivine crystals, once considered to be due to long-term annealing (Saiki et al., 2003), is now thought to have occurred primarily from resorption at high temperatures (above ~1250°C) in the presence of silicate melt (Boesenberg et al., 2012); a Thomson (Widmanstätten) structure was developed in the FeNi-metal component.
Boesenberg et al. (2012) propose a model in which the formation of a pallasite layer is the result of progressive fractional melting of a chondritic body. Over time, metal–silicate separation occurs producing an insulating crust and regolith along with a sulfide-rich metallic core, and in the molten outer core, buoyant olivine crystals coalesce and form a dunite layer. Heat from the molten core causes convection in the overlying mush of olivine+silicate melt+molten metal, wherein olivine near the dunite layer is entrained in the molten metal which promotes the downward crystallization of metal forming a pallasitic assemblage. A less significant role was inferred for impacts, by which shock waves produce fragmentation of some olivines; some fragmental olivines are subsequently rounded through partial resorption within the silicate melt.

In a subsequent experiments, Donohue et al. (2018) expanded upon this fractional melting model (see diagram below). They contend that as temperatures decrease from peak values of ~1600–1700°C to ~1000°C, at rates of ~100–300°C/m.y., minor phases crystallize through redox reactions and grain boundary diffusion. Partial equilibration occurred over a timescale of a few million years, altering the elemental distribution among the olivine, metal, and minor phases. Phosphorus from the molten metal was taken up into the residual silicate melt and ultimately formed phosphates (in a cooling sequence of merrillite, stanfieldite, farringtonite, Fe-rich phosphate, and silico-phosphate), with phosphoran olivine remaining as melt is exhausted. Orthopyroxene, chromite, and schreibersite were also formed as late-stage phases. Cooling rates gradually decreased to relatively slow rates of ~1°C/m.y. as determined through metallographic cooling models. standby for pallasite formation model diagram
click on image for a magnified view

Diagram credit: Donohue et al., GCA, vol. 222, p. 315 (2018)
‘Experimentally determined subsolidus metal-olivine element partitioning with applications to pallasites’

Scenario 2 (Hsu, 2003)

  1. Olivine crystallized as a fractionation cumulate from the silicate liquid in a magma chamber (as suggested by the lack of a trapped melt component, and consistent with elemental trends), or as a partial melt residue; ~50–70% melting is indicated
  2. A high-energy impact(s) resulted in the high-pressure injection of low-viscosity metal into the crystalline olivine layer.
  3. The pallasite material experienced very rapid cooling at high temperatures and slow cooling at low temperatures, consistent with the preservation of separate olivine and FeNi-metal and of zoning profiles (e.g., Ni) in the olivine.
  4. Evidence of live 53Mn, as well as other chronometric data, indicates that pallasites were formed within the first 10 m.y. of solar sytem history.
  5. A later event(s) produced an extensive regolith, which buried the pallasite material and initiated a period of slow cooling.
  6. Rounding of olivine crystals, once considered to be due to long-term annealing (Saiki et al., 2003), is now thought to have occurred primarily from resorption at high temperatures (above ~1250°C) in the presence of silicate melt (Boesenberg et al., 2012); a Thomson (Widmanstätten) structure was developed in the FeNi-metal component.
The olivine-metal mixing event could have resulted from the impact of a differentiated body having a fractionated liquid iron core onto another differentiated protoplanetary object very early in Solar System history—as early as ~4.557 b.y. ago and not later than ~4.3 b.y. ago. The injection of molten metal from the impactor created impact-melt, dike-like intrusions in the cold olivine mantle of the host body, forming a pallasitic mixture that was first rapidly frozen and then slowly cooled over a period of at least several tens of millions of years. Isotopic data suggest that this main-group pallasite parent body formed in the terrestrial planet-forming region. Thereafter, one or more severe impacts sent pallasitic fragments into parking orbits within the asteroid belt.

A study conducted by Tarduno et al. (2012) is most consistent with scenario 2. Paleointensity data were obtained by from sub-µm to µm-sized, stable magnetic inclusions within Imilac and Esquel olivines, continuing with analyses of Springwater (Tarduno and Cottrell, 2013). Along with cooling rate data, these inclusions indicate that pallasites formed and cooled under the influence of a strong magnetic field generated by a core dynamo on an ~320-km-diameter parent body. This remanent magnetization attests to the fact that the Imilac pallasite was not formed near the core-mantle boundary, because a rotating core dynamo would necessarily cease prior to any significant cooling of adjacent material; therefore, no remanent magnetization would exist. Their estimates of the cooling rate for pallasite material based on conduction (2–9K/m.y.) are consistent with a formation location within the upper ~60% of the protoplanet mantle—perhaps at depths of 10 km and 40 km for Imilac and Esquel, respectively.

Scenario 3 (Yang, 2010, Yang et al., 2010)

  1. Olivine crystallized from the silicate liquid at the lowest layer of the mantle, the core–mantle interface.
  2. The metallic core solidified outwards until ~80 vol% crystallization was reached.
  3. A glancing impact disrupted the protoplanet resulting in the high-pressure injection of the residual low-viscosity metal into the crystalline olivine layer from the lower mantle.
  4. Diverse cooling rates reflect cooling at different depths on a common parent bodyand not at the core–mantle interface.
  5. A rubble-pile asteroid was formed providing a source for main group pallasites.
  6. Rounding of olivine crystals, once considered to be due to long-term annealing (Saiki et al., 2003), is now thought to have occurred primarily from resorption at high temperatures (above ~1250°C) in the presence of silicate melt (Boesenberg et al., 2012).; a Thomson (Widmanstätten) structure was developed in FeNi-metal regions of appropriate size.
This scenario was the basis for the PSRD article by E. Scott, J. Goldstein, and J. Yang: ‘Formation of Stony-Iron Meteorites in Early Giant Impacts‘, June 2010. The diagram shown below demonstrates the general sequence of events proposed, initiated by a glancing impact between a differentiated body and a larger object, and culminating in the reassembly of the former into a much smaller pallasite parent body.

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In his study of main-group pallasites, Edward Scott (2017) expanded upon this formation model by inferring the existance of three zones at the core–mantle boundary (see diagram below). Initially, differentiation of the parent body occurred forming a molten metallic core and a dunitic olivine mantle. As olivine crystallized and accumulated at the base of the mantle, some portion became immersed in the molten iron where the crystal edges underwent rounding (zone 3 in the diagram); this rounding was once considered to be due to long-term annealing (Saiki et al., 2003), but is now thought to have occurred primarily from resorption at high temperatures (above ~1250°C) in the presence of silicate melt (Boesenberg et al., 2012). Ultimately, a glancing impact disrupted the parent body resulting in fragmentation of both the early-formed, Ir-rich (0.7–5 ppm) olivine located in zone 3, and the late-formed, Ir-poor (0.01–0.3 ppm) olivine located in zones 1 and 2. During the impact event, molten iron was injected into the olivine assemblages to produce large regions of pure metal within pallasitic zones.

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Image credit: Edward R. D. Scott, 48th LPSC, 2017, #1037

In their extensive elemental analysis of pallasites, Wasson and Choi (2003) proposed that gases associated with the metallic melt were concentrated in voids formed by core contraction and mantle collapse during cooling, and that subsequent condensation of these gases introduced enrichments of the volatile siderophiles Ge and Ga into the PMG members, as well as enrichments of Fa into the PMG-as members. They also attributed the refractory siderophile enrichments present in many pallasites (e.g., Ir) to the mixing of late-stage core metal and residual mantle metallic melts.

A study was conducted by Mittlefehldt and Herrin (2010) pertaining to the degree of magmatic fractionation of main-group pallasites, including anomalous members. They examined the Mn/Mg ratios of these pallasites and determined that there was no correlation between magmatic fractionation and metal composition. This realization was inconsistent with a core–mantle boundary origin of the olivine on a single parent asteroid. In a comparison of elemental abundances to the Mn/Mg ratios of the various pallasites, they found that the olivine was not formed through accumulation processes, but instead was formed as a residue of a high degree igneous melt.

The pallasite thermal history reflects a slow cooling rate of a few degrees per million years, as evidenced by the FeNi-metal component cooling over the temperature interval of ~700°C to ~500°C, which is the interval over which the Thomson (Widmanstätten) structure is formed (Lavrentjeva, 2009). This slow cooling rate is in contrast to the much more rapid cooling rate of a few degrees per year reflected in the olivine component at high temperature conditions of ~1100°C. The olivine diffusion gradients and other thermal history details are more consistent with an impact-generated mixture of core and mantle materials than a core–mantle boundary origin. Anomalous metal and silicate compositions measured in some pallasites might reflect solid–liquid metal mixing on a single main-group pallasite parent body consistent with common O-isotopic compositions on each. Radiometric dating indicates that such an impact occurred <10 m.y. after chondrule formation.

A novel hypothesis addressing pallasite formation was proposed by Asphaug et al. (2006), and was adapted by Danielson et al. (2009) to account for the wide variety of metal-silicate textures and bulk compositions observed in pallasites. They assert that pallasite diversity could be attributed to their formation on a chain of objects that was produced as a result of a grazing collision between partially molten Moon- to Mars-sized planetary embryos. These may be represented by multiple disparate pallasite groups such as (i) Brenham, (ii) Imilac, (iii) Fukang, and (iv) Seymchan (Johnson et al., 2010). Uniquely similar volatile element depletions that exist between the pallasites and the HED meteorites suggest a possible association between these different planetary bodies. These facts prompt speculation that these two planetesimals, while in their embryonic stages early in Solar System history, experienced a mutual grazing collision.

At least as intriguing is a formation hypothesis envisioned by M. Fries (2012) in which pallasites formed in the cores of small, spherical, rapidly cooled bodies in which gravitational differentiation is at a minimum and convective forces are insignificant. Such quiescent conditions would allow silicates to remain in the core while molten metal slowly infiltrated and disaggregated the silicates into ever smaller angular fragments. A subsequent catastrophic impact disruption of the parent body sent portions of this pallasitic core into Earth-crossing orbits.

The metal and O-isotopic compositions of the main-group pallasites, including the phosphoran nature of olivine in some members (Brahin, Brenham, Rawlinna 001, Springwater, and Zaisho), are consistent with features of late-stage crystallization (high-Au, ~80% core crystallization) of residual melts in the IIIAB iron core. However, recent studies appear to rule out a genetic connection to IIIAB irons and a core–mantle boundary formation scenario (Yang and Goldstein, 2006; Yang et al., 2010). New and more precise metallographic cooling rates were obtained for pallasites utilizing taenite Ni compositions, cloudy zone particle sizes, and tetrataenite bandwidths, the latter two parameters being positively correlated with each other and negatively correlated with the metallographic cooling rates derived from taenite. The results are not what one would expect given an origin at the core–mantle boundary. Instead, based on the size of the taenite particles (island phase) in the cloudy zone of the pallasites, as well as on the tetrataenite bandwidth, the cooling rates were demonstrated to have a wide range inconsistent with a core–mantle boundary of a solitary asteroid. Cooling rates were significantly lower for pallasites than for IIIAB irons, with rates of 2.5–18K/m.y. measured for main-group members and 13–16K/m.y. measured for the Eagle Station group, while IIIAB irons cooled at ~50–350K/m.y. This implies that the irons were actually closer to the surface than the pallasites. Paradoxically, the ungrouped pallasite Milton, which lacks cloudy taenite zones and did not experience shock reheating, exhibits a cooling rate >5000K/m.y. (Yang et al., 2010).

In their measurement of high-Ni particles within the cloudy zone of several main-group pallasites and IIIAB irons, Yang et al. (2007) found that a correlation exists between cooling rates and bulk Ni in IIIAB irons but not in main-group pallasites. Based on the significantly larger size of the high-Ni metal particles in pallasites (82–170 nm) than in the IIIAB irons (42–58 nm), they determined that the cooling rate was ~2.5–25 times slower in the pallasites, with the wide range of cooling rates indicative of a large thermal heterogeneity within the pallasite formation zone which did not exist on the IIIAB iron parent body. Notably, the Re–Os chronometer suggests that pallasites formed 60 m.y. later than IIIAB irons, raising further doubt about a IIIAB core–mantle origin for main-group pallasites (E. Scott, 2007).

Further evidence in support of separate parent bodies for main-group pallasites and IIIAB irons was provided by Huber et al. (2011). They found that pallasites have a much younger range of cosmic ray exposure (CRE) ages than the IIIAB irons. In an effort to better resolve the CRE age difference between main-group pallasites and IIIAB irons, Herzog et al. (2015) conducted highly precise cosmogenic radionuclide analyses of both metal and olivine components in a large number of main-group pallasites. Utilizing multiple dating systems, they demonstrated that a significant number of these pallasites define a broad cluster of ages near 100 m.y., while only a very few of the IIIAB irons measured (6 of 33; Herzog and Caffee, 2014) fall within this range—most members of this iron group have much older CRE ages. They concluded that at least half of the main-group pallasites are associated with just a few common ejection events on their parent body, and that the IIIAB irons probably derive from a separate parent body.

Previous O-isotopic analyses for main-group pallasites and the HED meteorites indicated that these two groups have values that are very similar. In a high precision comparative analysis of the oxygen three-isotope composition between olivines from five main-group pallasites and representative HED samples, including eucrite and diogenite material, Jabeen et al. (2013) determined that a clear distinction exists, thus demonstrating that these meteorite groups originated on separate parent bodies. In another study investigating the close O-isotopic relationship between main-group pallasites, mesosiderites, and the HED clan, Ziegler and Young (2007) discovered that non-homogenized samples of main-group pallasite olivines exhibit a bimodality in 17O values, which also distinguishes their origin from that of the mesosiderites and the HED clan. In a follow-up study, a more refined O-isotopic analysis was conducted by Greenwood et al. (2008), but their results did not support a bimodality in 17O values; however, they definitively established that the parental source of main-group pallasites was different from that of mesosiderites and the HED clan.

Subsequent high-precision triple oxygen isotope analyses of a broad sampling of main-group pallasites (Brahin, Brenham, Esquel, Fukang, Giroux, Huckitta, Imilac, Seymchan, Springwater, and Sterley) and selected members of the HED group (Tatahouine, Stannern, and Juvinas) were conducted by Ali et al. (2013, 2014). Their results, together with geochemical and other data, not only demonstrate that the HEDs are not genetically related to the main-group pallasites, but also that a bimodality exists for these pallasites based on several factors: Δ17O values, MgO content in olivines, bulk olivine abundance, concentration density of olivine grains, and paleointensity. They were able to resolve systematic variations among the main-group pallasites which indicate the existence of two distinct subgroups (see diagram below). This O-isotopic bimodality has been attributed to several possible scenarios, including the existance of multiple parent bodies, the sampling of different locations on a common parent body, and/or varibility in the degree of impactor contamination.

  1. high-Δ17O (ave. –0.172 [±0.007] ‰), less olivine-rich (olivine/metal = 2.0); e.g., Brenham, Huckitta, Imilac, Springwater, Sterley
  2. low-Δ17O (ave. –0.213 [±0.011] ‰), more olivine-rich (olivine/metal = 2.9); e.g., Brahin, Esquel, Fukang, Giroux, Seymchan

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Diagram credit: Ali et al., 45th LPSC, #2390 (2014) Another high-precision oxygen isotope analysis was undertaken by Greenwood et al. (2015) in which 24 main-group pallasites (Admire, Ahumada, Brahin, Brenham, Dora, Esquel, Finmarken, Fukang, Giroux, Glorietta Mountains, Imilac, Krasnojarsk, Lipovsky, Marburg, Marjalahti, Molong, Pallasovka, Pavlodar, Quijingue, Rawlinna 001, Santa Rosalia, Somervell County, Springwater, and Theil Mountains) and a number of mesosiderite olivine-rich clasts and related dunites (Lamont, Mount Padbury, Vaca Muerta, NWA 2968, NWA 3329) were utilized. Their results support the previous findings showing that the main-group pallasites and HED meteorites originated on separate parent bodies. However, the new Δ17O values of the 24 main-group pallasites studied do not support the previous hypothesis for bimodality, but instead indicate that a continuum exists having an average Δ17O value of –0.187 (±0.016) ‰.

More recently, Ali et al. (2018) employed improved laser fluorination techniques to increase the precision of triple oxygen isotope data for 25 MG pallasites. They determined that a significant bimodality exists, and it is clearly demonstrated that two statistically distinct subgroups are resolved. These subgroups likely represent at least two asteroidal parent bodies with each having homogeneous olivine compositions.

  1. high-Δ17O-bearing (ave. –0.166 [±0.014] ‰) subgroup; e.g., Acomita, Ahumada, Brenham, Finmarken, Huckitta, Imilac, Jay Bird Springs, La’gad 002, Marjahlati, Otinapa, Pallasovka, Somervell County, South Bend, Springwater, Sterley, Thumrayt 001
  2. low-Δ17O-bearing (ave. –0.220 [±0.009] ‰) subgroup; e.g., Brahin, Esquel, Fukang, Giroux, Hambleton, Krasnojarsk, Mount Dyrring, Newport, Seymchan

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Diagram credit: Ali et al., MAPS, vol. 53, #6, p. 1228 (2018)
‘The oxygen isotope compositions of olivine in main group (MG) pallasites: New measurements by adopting an improved laser fluorination approach’
Utilizing the paleointensity data of Tarduno et al. (2012) for the low-Δ17O Esquel and the high-Δ17O Imilac, Ali et al. (2014) ascertained that they each formed at different depths (40 km and 10 km, respectively) on one or more parent bodies. Paleointensity data was compiled by Nichols et al. (2018) for five pallasites representing a wide range of cooling rates (3–8°C/m.y.). These data were used to demonstrate the evolution of a late-stage core dynamo on the parent body beginning ~100 m.y. after accretion and spanning a period of ~140 m.y. (see diagram below showing relative paleointensities, where Imilac = 73.6 [±8.1] µT as determined by Tarduno et al., 2012). standby for pallasite paleointensity diagram
Diagram credit: Nichols et al., 49th LPSC, #1976 (2018) The question pertaining to whether a genetic relationship exists among HED-clan meteorites, mesosiderites, main-group pallasites, and IIIAB irons is ongoing. It is now recognized (e.g., Sanborn et al., 2014) that a comparison of Δ17O vs. ε54Cr is one of the best diagnostic tools for determining genetic relationships between meteorite groups. Moreover, Sanborn et al. (2015) demonstrated that ε54Cr values are not affected by aqueous alteration. Utilizing both the ε54Cr and Δ17O values for representative samples of each of these meteorite groups, Wasson and Göpel (2014) found that these groups were unresolvable in terms of ε54Cr values, and that the differences in Δ17O values are reasonable given a scenario of rapid impact-heating for the HED meteorites. They argue that isotopic evidence which supports an origin of the HED meteorites on the IIIAB parent body should be considered more reliable than any association of the HED-clan meteorites with asteroid 4 Vesta based on spectral analyses from orbit. standby for 54Cr vs 17O diagram photo
Diagram credit: Wasson and Göpel, 77th MetSoc, #5446 (2014) In an effort to better resolve potential genetic relationships that might exist among meteorite groups, a Cr-isotopic analysis was conducted by Sanborn et al. (2018) for olivine from both the main-group pallasite Brenham and the ungrouped pallasite Milton, along with the anomalous IVA irons Steinbach and São João Nepomuceno. It is demonstrated on a coupled Δ17O vs. ε54Cr diagram (shown below) that Brenham and Krasnojarsk plot significantly above the HED normal trend (black squares in inset), which supports the inference that these meteorites formed on separate parent bodies. Chromium vs. Oxygen-isotope Plot
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click on diagram for a magnified view

Diagram credit: Sanborn et al., 49th LPSC, #1780 (2018) It has been proposed that the solid inner core of the main-group pallasite parent body measured up to 950 km in diameter, and that it was 80% solidified at the time it was separated from the remaining 20% melt during a glancing collision with a larger body (Yang et al., 2010). The Ir-poor residual melt was then mixed with twice the volume of olivine mantle fragments to form a body up to 800 km in diameter (smaller with a silicate regolith). Utilizing temperature and pressure constraints for the stability of tridymite inclusions present in the Fukang pallasite, Della-Giustina et al. (2011) determined the maximum size limit for the main-group pallasite parent body to be ~600 km in diameter; a minimum size still large enough to enable differentiation would be ~40 km. Further modeling was conducted by Habib et al. (2018) to better constrain the size of the main-group pallasite parent body. Employing mass and pressure data as a function of the parent body radius, and interpreted under the assumption of a core–mantle boundary origin for the Fukang pallasite, they calculated a maximum diameter of 690–1350 km for the body. To date no iron meteorites have been found which originated on the main-group pallasite parent body, suggesting that little olivine-free metal survived the collision.

Based on all of the data gathered so far, it could be concluded that the pallasites in our collections represent at least seven separate parent bodies: 1) main-group; 2) Eagle Station group; 3) Milton; 4) Choteau + Vermillion + Y-8451; 5) Zinder + NWA 1911; 6) NWA 10019; 7) LoV 263. In addition, several pallasites with anomalous silicates (e.g., Springwater) and anomalous metal (e.g., Glorieta Mountain) could possibly increase the number of unique parent bodies. The Imilac specimen pictured above is a 56.0 g quarter slice (lower left quadrant) sectioned from the 1.57 kg mass shown in the top photo below. The bottom photo is a large slice showing the typical distribution of silicate and metal in Imilac, courtesy of Sergey Vasiliev. standby for imilac photo
Photo courtesy of Alan Lang—R.A. Langheinrich Meteorites

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Photo courtesy of Sergey Vasiliev—

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