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.
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 can occur twice as large. The metal in the smaller specimens shows evidence of violent shearing and deformation with frictional heating reaching recrystallization temperatures. As in most pallasites, 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 silicatemetal mixing. The second requires impact-induced injection of molten metal into olivine. The third also presupposes the injection of molten metal into olivine, but at a near-surface location.
olivine crystallized from the silicate liquid at the lowest layer of the mantle, the coremantle interface.
cooling and contraction of the metallic core produced a 2% volume reduction leaving a void at the coremantle boundary.
the overlying crystalline olivine then collapsed into the viscous metal where heating and mixing occurred to produce the pallasitic structure.
over an extended cooling period, grain boundary energy was reduced through annealing of olivine molecules within the lattice, thus producing a rounding of olivine crystals (Saiki et al., 2003); a Thomson (Widmanstätten) structure also developed.
Scenario 2 (Yang, 2010, Yang et al., 2010)
olivine crystallized from the silicate liquid at the lowest layer of the mantle, the coremantle interface.
the metallic core solidified outwards until ~80 vol% crystallization was reached.
a glancing impact(s) disrupted the protoplanet resulting in the high-pressure injection of the residual low-viscosity metal into the crystalline olivine layer from the lower mantle.
Diverse cooling rates reflect cooling at different depths on a common parent bodyand not at the coremantle interface.
a rubble-pile asteroid was formed providing a source for main group pallasites.
over an extended cooling period, following olivine fragmentation, grain boundary energy was reduced through volume diffusion of olivine molecules within the lattice (annealing), thus producing a rounding of olivine crystals (Saiki et al., 2003); a Thomson (Widmanstätten) structure also developed in FeNi-metal regions of appropriate size.
Scenario 3 (Hsu, 2003)
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; ~5070% melting is indicated
a high-energy impact(s) resulted in the high-pressure injection of low-viscosity metal into the crystalline olivine layer.
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.
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.
a later event(s) produced an extensive regolith, which buried the pallasite material and initiated a period of slow cooling.
over time, annealing produced a rounding of olivine crystals (Saiki et al., 2003), while a Thomson (Widmanstätten) structure was developed in the FeNi-metal component.
A more recent study is most consistent with scenario three. Paleointensity data was obtained by Tarduno et al. (2012) 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 ~160-km-radius parent body. Their estimates of the cooling rate for pallasite material based on conduction (29°K/m.y.) are consistent with a formation location within the upper ~60% of the protoplanet mantleperhaps at depths of 10 km and 40 km for Imilac and Esquel, respectively. The olivine-metal mixing event likely resulted from the impact of a differentiated body having a fractionated liquid iron core onto another differentiated protoplanetary object very early in solar system historyas 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.
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 coremantle 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 metal and O-isotopic compositions of the main-group pallasites, including the phosphoran nature of olivine in some members (Springwater, Brahin, Rawlinna, 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 coremantle 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 coremantle 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 coremantle boundary of a solitary asteroid. Cooling rates were significantly lower for pallasites than for IIIAB irons, with rates of 2.518°K/m.y. measured for main-group members and 1316°K/m.y. measured for the Eagle Station grouplet, while IIIAB irons cooled at ~50350°K/m.y. This implies that the irons were actually closer to the surface than the pallasites were. Paradoxically, the ungrouped pallasite Milton, which lacks cloudy taenite zones and did not experience shock reheating, exhibits a cooling rate >5000°K/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 (82170 nm) than in the IIIAB irons (4258 nm), they determined that the cooling rate was ~2.525 times slower in the pallasites, with the wide range of cooling rates indicative of a large thermal heterogeneity within the pallasite formation zone, which was not on the IIIAB iron parent body. Notably, the ReOs chronometer suggests that pallasites formed 60 m.y. later than IIIAB irons, raising further doubts about a IIIAB coremantle origin for main-group pallasites (E. Scott, 2007). Equally important, pallasites have a much younger range of CRE ages than the IIIAB irons (Huber et al., 2011).
The pallasite thermal history reflects a slow cooling rate of a few degrees per million years, exhibited in the FeNi-metal component over the temperature interval of ~700°C to ~500°C, 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 coremantle boundary origin. Anomalous metal and silicate compositions measured in some pallasites may reflect solidliquid 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 may 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.
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 their follow-up of this report, 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-resolution oxygen three-isotope analyses of a broad sampling of main-group pallasites were conducted by Ali et al. (2013). Their results, together with geochemical and other data, demonstrate that a bimodality does exist 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 in their study that suggest the existence of two distinct subgroups as follows:
1. low-Δ17O (e.g., Esquel, Seymchan, Brahin, Fukang, Giroux)
2. high-Δ17O (e.g., Imilac, Brenham, Springwater, Huckitta, Sterley)
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.
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.
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). However, utilizing the pressure constraints on the tridymite inclusions present in the Fukang pallasite, the maximum size limit for the main-group pallasite parent body would be ~600 km in diameter; a minimum size still large enough to enable differentiation would be ~40 km (Della-Giustina et al., 2011). 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. The Imilac specimen pictured above is a 56.0 g quarter slice sectioned from the 1.57 kg mass shown in the top photo below (lower left quadrant). The bottom photo is a large slice showing the typical distribution of silicate and metal in Imilac, courtesy of Sergey Vasiliev.