The Conventional View of Chondrites
Chondrites have conventionally been interpreted as aggregates of primitive materials that were assembled at the very start of the solar system from the same reservoir of nebular dust as went to make the Sun (e.g., Wood 1988). This view stems from the remarkable similarity between the chemical composition of chondrites and that of the Sun’s photosphere for all elements other than a few that normally occur in gases (e.g., H, He, C, N, Ar). The view is reinforced by the presence in chondrites of CAIs, which are the oldest dated objects with a solar system isotopic signature (Amelin et al. 2002, 2010). Indeed, the time of CAI formation has now been widely adopted as defining the start of the solar system (t = 0). In addition, chondrites contain pristine grains of stardust that are even older than the solar system (e.g., Hoppe 2008). Thus, conventional thinking holds that chondrules, along with CAIs, were created directly from clumps of nebular dust at the outset, before the first planetesimals (presumed to be the chondrite parent bodies) had accreted. It further holds that after their accretion, some chondritic planetesimals became overheated, melted, and differentiated into molten metal cores and basaltic crusts, the sources, respectively, of iron and basaltic meteorites (e.g., Lauretta and McSween 2006).
However, this intuitive and long-established interpretation of meteorites has recently been challenged by chronological evidence, which suggests that chondrules were made after, and not before, the parent bodies of differentiated meteorites had melted.
The Chronology of Molten Cores and Chondrules
Most iron meteorites now appear to come from planetesimals that had accreted and melted extremely early, perhaps within 1 Myr of CAI formation (t < 1 Myr) because their ε182W values are very low and within error of the initial ε182W of CAIs (Burkhardt et al. 2008). In the early solar system, ε182W was rising due to the radioactive decay of 182Hf to 182W (half-life 8.9 Myr). Tungsten is a siderophile (iron-loving) element, whereas hafnium is lithophile (silicate-loving). The ε182W values of iron meteorites became fixed, therefore, when molten metal segregated into planetesimal cores. At that stage, the radioactive hafnium, being lithophile, was removed from close proximity to metal and transferred to the silicate mantle of each planetesimal where subsequent radiogenic 182W accumulated. A recent estimate of the time of core formation based on ε182W in magmatic iron meteorites is 0.3 ± 1.2 Myr after CAIs (Kruijer et al. 2011). Even more recently, Burkhardt et al. (2012) revised the value of initial ε182W in CAIs, and they inferred that while the cores of some parent bodies (notably the IVB group) formed at t approximately 0.3 Myr, the cores of others continued to form up to about t = 2 Myr. The timing of core separation is critical to constraining early disk evolution, and is the culmination of painstaking efforts to understand and refine the Hf-W chronometer by many workers (Horan et al. 1998; Kleine et al. 2005; Markowski et al. 2006; Scherstén et al. 2006; Qin et al. 2008; Burkhardt et al. 2008). In particular, Markowski et al. (2006) recognized the need to correct ε182W values for the effects of long-term exposure to cosmic rays; uncorrected values had previously given erroneously old ages.
By contrast, 26Al-26Mg dating suggests that chondrules are mostly 1.5–2.5 Myr younger than CAIs. Following the pioneering work of Hutcheon and Hutchison (1989), about 100 chondrule dates based on 26Al-26Mg internal isochrons have now been published. A recent review of them by Kita and Ushikubo (2012) shows that more than 90% of those from unequilibrated (type 3.0) chondrites (64 determinations) fall within the 1.5–2.5 Myr age range, with just three dating from t < 1.5 Myr. The same age difference between chondrules and CAIs of about 2 Myr has been measured independently by 207Pb-206Pb dating (e.g., Amelin et al. 2002, 2010; Connelly et al. 2008), and by 182Hf-182W dating (Kleine et al. 2008). As the chondrite parent asteroids were assembled after the youngest chondrules within them had formed, i.e., probably more than about 2.5 Myr after CAIs, then far from being the first bodies to have accreted, as is conventionally assumed, they were perhaps among the last to have done so.
The late accretion of chondritic asteroids was already suspected on petrographic grounds long before the recent chronological evidence became known. Fragments deemed to be of planetary igneous rock were identified in chondrites by Kurat and Kracher (1980), Hutchison et al. (1988), and Kennedy et al. (1992). The last of these authors reported a 2 mm chip of high Mn/Fe basalt in the Parnallee chondrite, which they interpreted as being derived from a high Mn/Fe planetesimal that had already melted and broken up before Parnallee’s parent body had accreted. In addition, Ruzicka et al. (1995) reported an unusual silica pyroxenite clast in Bovedy (L3), which they regarded as having a planetary igneous origin. More recently, Sokol et al. (2007) reviewed the occurrence of a wide variety of differentiated igneous rock fragments in chondrites, and Ruzicka et al. (2012b) reported further silica-rich clasts of supposed igneous origin. On a related note, Libourel and Krot (2007) discovered small pieces of texturally equilibrated olivine rock inside chondrules, which they interpreted as tiny fragments of earlier planetesimals that had been metamorphosed and then broken up by impacts before being incorporated into chondrules. Although Whattam et al. (2008) questioned that interpretation, the petrographic evidence, like the chronological evidence, clearly points to high-temperature planetesimal processing prior to chondrite accretion.
Heating by 26Al: The Key to the New Chronology
The cause of early melting and core formation, some 1–2 Myr before chondrules were made, is not hard to find. Ever since evidence for live 26Al was discovered in CAIs, it has been realized that the decay energy from this short-lived isotope (half-life 0.72 Myr) would have been more than sufficient to melt the fully insulated interiors of planetary bodies that accreted early enough, while radioactive heating was intense (Lee et al. 1977). The corollary is that the chondrite parent bodies, as they did not melt, accreted later, after the 26Al had largely decayed and lost its capacity to cause melting. This explanation is reinforced in the following paragraphs by a simple quantitative analysis of the likely effects of 26Al heating on the timing of initial planetesimal meltdown, the timing of chondrule formation, and the timing of chondrite accretion.
Our estimate of the energy available in 26Al to heat the first crop of planetesimals at t = 0 is about 6.6 kJ per gram of dry dust. This estimate requires knowledge of the concentration of Al in the dust, of the initial ratio of 26Al/27Al, and of the heat released by each decaying atom of 26Al. The concentration of Al is not known precisely, but would presumably have been more than 0.85 wt.% (the level in CI chondrites, which are extensively hydrated; Lodders and Palme 2009). We conservatively, although somewhat arbitrarily, choose a value of 1.2 wt.%, which corresponds roughly to the concentration of Al in dehydrated CI chondrite, and is close to the concentration of Al in most anhydrous chondrite groups (Lodders and Fegley 1998). We assume that the initial value of 26Al/27Al in the disk was uniformly 5 × 10−5, the so-called canonical value in CAIs (Jacobsen et al. 2008; MacPherson et al. 2010). A uniform distribution of canonical 26Al in the disk is indicated by the identical 26Mg/24Mg in the Earth, the Moon, Mars, and bulk chondrites (Thrane et al. 2006); by the correlation between the initial 26Mg/24Mg and the 26Al/26Mg ages of a suite of chondrules studied by Villeneuve et al. (2009); and by time intervals between specific events measured using the 26Al/26Mg chronometer being corroborated by other chronometers (e.g., Connelly et al. 2008). We are aware that Larsen et al. (2011) reported significant variation in 26Mg/24Mg in objects with solar 27Al/24Mg, and proposed that the initial 26Al/27Al in parts of the inner solar system where planetesimals accreted and where chondrules formed may have been substantially lower than the canonical level where CAIs were made. However, Wasserburg et al. (2011) found a wide variation in the initial 26Mg/24Mg of different CAIs with identical canonical 26Al/27Al, which leaves an open verdict for the case made by Larsen et al. (2011). Finally, we take the decay energy per atom of 26Al as 3.1 MeV (Castillo-Rogez et al. 2009). A plausible 10% uncertainty in both the initial 26Al/27Al, and in the wt. % Al, leaves our estimated initial radioactive energy at 6.6 ± 1 kJ g−1.
In addition to 26Al, the short-lived isotope 60Fe may have contributed to radioactive heating. However, the initial concentration of 60Fe and whether it was uniformly distributed in the disk remain unknown (Telus et al. 2012). The ratio at t = 0 of 60Fe/56Fe (1.5 × 10−6) assumed by Sanders and Taylor (2005) now seems far too high. Telus et al. (2011) suggest that it was between 3 and 5 × 10−7, making the contribution of 60Fe to heating <0.5 kJ g−1. Moreover, as 60Fe’s half-life of 2.6 Myr (Rugel et al. 2009) is more than three times longer than that of 26Al (0.72 Myr), its contribution to overall heating during the critical first 2 Myr would have been trivial, and we therefore ignore it in this paper.
Figure 1a shows the temporal decline in energy stored as 26Al in each gram of dry primitive dust, starting from the initial 6.6 kJ g−1, through almost seven half-lives during the first 5 Myr. To put this decline in perspective, 6.6 kJ g−1 is about four times larger than the 1.6 kJ g−1 needed to fully melt the insulated interior of a planetesimal at a temperature of 1850 K. The estimate of 1.6 kJ g−1 assumes starting from cold (250 K), with specific heat capacity, Cp = 837 J kg−1 K−1, and latent heat of fusion = 2.56 × 105 J kg−1 (Hevey and Sanders 2006). Thus, planetesimals that accreted during the first two half-lives of 26Al, or roughly during the first 1.5 Myr, would have had the potential to become completely molten in their fully insulated interiors.
Figure 1. a) Exponential decline with time of the potential thermal energy stored as 26Al in a gram of “dry” primitive dust. b) Time at which solidus (1425 K) and liquidus (1850 K) temperatures are reached in the fully insulated interior (deeper than approximately 5 km at t = 1 Myr to deeper than approximately 20 km at t approximately 5 Myr—see Figs. 2 and 4) of a planetesimal as a function of the time of its cold (250 K) instantaneous accretion and assuming no melt migration during heating. Arrows A, B, C, and D illustrate the timing of initial and total melting following cold accretion at t = 0, 0.75, 1.4, and approximately 2 Myr, respectively (see text for explanation). Accretion-time intervals labeled 1, 2, 3, and 4 relate to the fields shown in Fig. 4. The lower edge of the gray zone is the 1850 K liquidus calculated using latent heat and specific heat capacity values from Ghosh and McSween (1999), which are greater than those adopted here.
Download figure to PowerPoint
Figure 1b shows the time it would have taken to reach the onset of melting (the solidus temperature, approximately 1425 K) and also the completion of melting (the liquidus temperature, approximately 1850 K) of the fully insulated interior (i.e., with zero heat loss) as a function of the time of cold planetesimal accretion, assuming no migration of the 26Al heat source. With accretion at t = 0 (arrow “A” in Fig. 1b), heating would have been rapid and the liquidus would have been reached by about t = 0.3 Myr. This is in good agreement with the timing of earliest planetesimal melting and core formation, shown by 182W-deficit dating of iron meteorites (Burkhardt et al. 2008, 2012; Kruijer et al. 2011) and, although errors in the dating are large, such early melting clearly endorses the assumption that 26Al was the heat source.
With accretion at t = 0.75 Myr (arrow “B”), the initial heating rate would have been half that for arrow “A,” but rapid enough for total internal melting to have been achieved by t = 1.5 Myr.
As a third example, with accretion at t approximately 1.5 Myr (arrow “C”), the insulated interior of a planetesimal would have carried just enough 26Al to reach the liquidus, but melting would not have been completed until after t approximately 5 Myr. This example may explain the paucity of chondrules that date from before t approximately 1.5 Myr (Kita and Ushikubo 2012). Assuming that chondrules (regardless of their formation mechanism) were produced in large numbers before t approximately 1.5 Myr, the scarcity of those old chondrules in meteorites must reflect their poor survival rate. We imagine that such chondrules accreted to planetesimals before t approximately 1.5 Myr and became buried in their insulated interiors where they would later have been melted down and destroyed. If this explanation is correct, then it implies that these pre-1.5 Myr chondrules, once made, did not linger in space, but accreted to planetesimals promptly and were thence destined for a magmatic grave.
Finally, with accretion after about t approximately 2 Myr (arrow “D”), the level of 26Al would have been too low to have heated the planetesimal’s interior to the solidus, so no melting at all would have taken place. The timing is consistent with the evidence that chondrites (which of course did not melt) accreted after about t approximately 2.5 Myr (by when chondrule production was in decline), and again corroborates the view that 26Al was the main heat source within planetesimals.
In summary, the timing of core formation before t = 1 Myr, the scarcity of chondrules made before t = 1.5 Myr, and the accretion of chondrites after t = 2.5 Myr, combine to uphold our conviction that planetesimal heating by 26Al was a key factor in the evolution of planetesimals in the infant solar system.
The Structure of Molten Planetesimals
To visualize the changing internal structure of a molten planetesimal, we use the results of Hevey and Sanders (2006) who presented simulations of the heating, melting, and cooling of initially cold, porous planetesimals that accreted instantaneously. Their model is based on a radiogenic heat budget of 6.4 kJ g−1 of dust at t = 0, which is very close to the value of 6.6 kJ g−1 we estimate here. We note that they used an incorrect 26Al decay energy (4 MeV). That value wrongly includes approximately 1 MeV of energy that is not deposited as heat, but is lost in neutrinos. However, the error was fortuitously compensated by a lower concentration of Al (0.9 wt.%) compared with the 1.2 wt.% we use here.
As an example of their results, Fig. 2a illustrates the changing temperature profile within a planetesimal that accreted cold (250 K) at t = 0 and had a radius of 50 km (after early sintering and shrinkage). After a little over 0.3 Myr of heating, with the 26Al heat source evenly distributed, the interior deeper than approximately 5 km would have become uniformly hot and 50% molten (approximately 1725 K). At this stage, the interior is assumed to have lost rigidity and become cohesionless magmatic slurry undergoing turbulent thermal convection. With continued intense heating beyond 0.3 Myr, the magma is assumed to have remained at sub-liquidus temperatures, but to have increased in volume as the overlying rigid carapace was melted upward from its base and its thickness reduced from approximately 5 km at t = 0.3 Myr to just 0.5 km by t = 0.5 Myr (Fig. 2b). By that time, the rate of conductive heat loss through the residual crust would have reached a maximum, equaling the rate of internal heat production, and the crust’s thickness would have been at a minimum. Thereafter, with heat production lower than heat loss, no further melting would have occurred, and the crust would have thickened, slowly at first but ever more rapidly, over the next 2 or 3 Myr and beyond. A cartoon of the state of the planetesimal at t = 2 Myr is shown in Fig. 3.
Figure 2. a) Temperature profiles at selected times inside a planetesimal with a 50 km radius and zero porosity that accreted at t = 0 and a temperature of 250 K and was heated by 26Al decay. Broken lines are profiles during heating (until 0.5 Myr) and continuous lines are profiles during cooling. Convection began soon after t = 0.3 Myr, and by t = 0.5 Myr the molten, convecting interior had expanded to within about 0.5 km of the surface (after Hevey and Sanders 2006). b) Depth of solid rock and crust (gray) for the same body as a function of time.
Download figure to PowerPoint
Figure 3. Cartoon showing the internal structure of the 50 km radius planetesimal exemplified in Fig. 2 at t = 2 Myr. The core is assumed to be fully formed, but it is possible that small droplets of metal may have been held in suspension by turbulent convection (symbolized by curved arrows) in the magma ocean. The base of the crust is arbitrarily taken as the level at which the interior is 50% molten; the lower crust, with less than 50% melting, is deemed to be rigid. The 2 m of dust shown on the surface is predicted by instantaneous accretion; in reality, continuous accretion probably led to a considerable thickness of cool, loose, dusty debris, particularly after about t = 2 Myr when 26Al heating was very weak.
Download figure to PowerPoint
The model predicts that during maximum heat loss, only about 2 m of porous, unconsolidated, and extremely insulating dust separated solid, sintered rock at 700 K from the surface at 250 K. However, this prediction assumes that all accretion was completed instantaneously at t = 0. In reality, at least some accretion would have continued after the initial aggregation of material. Beyond about t approximately 2 Myr, with the 26Al heat source fading, any such late accretion would not have melted, but accumulated as a coating of loose, or weakly consolidated, dusty debris, which could have attained a considerable thickness.
What if the radius had been much smaller than the 50 km chosen in Fig. 2? Hevey (2001) showed that a body with a 20 km radius would have melted substantially by t = 0.3 Myr, and its crust would have thinned down to a minimum of 1.5 km by t = 0.5 Myr, but its high surface-to-volume ratio would then have led to rapid cooling, and the body would have been largely solid by t = 3 Myr. A body 10 km in radius would scarcely have melted at all. In a comparable thermal model, Moskovitz and Gaidos (2011) predicted similar melting behavior.
What if the radius had been larger? Going up in size, if the radius had been doubled, and was 100 km instead 50 km, the ratio of internal heat production to surface area would also have doubled, giving twice the heat flow and halving the thickness of insulating crust to a mere 250 m during the period of peak heat loss between t = 0.5 and t = 1.5 Myr. In this case, the crust may perhaps easily have foundered, exposing incandescent magma at the surface. With a still larger radius, the insulating carapace would have become even thinner and even more susceptible to foundering.
Figure 4 shows the effects of a planetesimal’s radius on its predicted melting behavior combined with the effects of the timing of its accretion discussed above (Fig. 1b). Four fields, numbered (1) to (4), correspond to the four accretion time intervals shown in Fig. 1b. Planetesimals starting in field (1) would have become extensively molten before t = 1.5 Myr, with very thin crusts as depicted in Figs. 2 and 3. Such planetesimals, we suggest below, would have been potential sources of chondrules by impact splashing. Planetesimals starting in field (2) would also have undergone extensive internal melting, but beneath a thicker insulating crust than for field (1), and with a longer heating period, generally becoming molten from t approximately 1.5 Myr up to t approximately 5 Myr depending on the time of accretion. Planetesimals starting in field (3) would have melted only partially, and they probably would have remained rigid. It is possible that the melt fraction (basalt magma) would have migrated upward, and that these planetesimals included the parent asteroids for primitive achondrites like the lodranites and ureilites. Planetesimals from field (4) would never have melted, although they may have become heated and metamorphosed. They would have become chondrite parent bodies. Figure 4 can be regarded as a refinement of the related, but oversimplified and rather misleading, two-field diagram presented by Hevey and Sanders (2006, fig. 6) in which planetesimals were shown either to have melted or not melted. It also bears similarities to fig. 5 of Moskovitz and Gaidos (2011), although the latter has later accretion times for given outcomes because it assumes 4 MeV, and not 3.1 MeV, as the decay energy per atom of 26Al.
Figure 4. Plot showing the eventual outcome of heating by 26Al decay in planetesimals as a function of radius and time of instantaneous cold accretion. The boundaries that delineate the four different fields are interpolated from Hevey and Sanders (2006, fig. 6) and Fig. 1b. Field (1) delimits planetesimals that will become substantially molten beneath a thin (e.g., <1 km) insulating crust before t approximately 1.5 Myr. Such planetesimals, we argue, will be prime candidates for bursting into chondrule spray if disrupted by impact between t approximately 1.5 and t approximately 2.5 Myr.
Download figure to PowerPoint
The thermal model of Hevey and Sanders (2006) assumes that the internal 26Al heat source remains evenly distributed at all stages during heating and cooling. However, some authors question this assumption, arguing that basalt migrates rapidly upward as soon as it is generated (e.g., Moskovitz and Gaidos 2011). As nearly all aluminum enters the basaltic melt fraction, the removal of such melt would also remove the heat source, forestall further internal heating, and invalidate the pattern of melting shown in Fig. 2. Wilson and Goodrich (2012) even suggested that basalt migration away from the zone of partial melting was so rapid that “high degrees of mantle melting never occurred in any asteroids.”
While we acknowledge that basalt was removed in the case of the ureilite parent body, which we believe to have accreted late, in field 3 of Fig. 4, we cannot accept that basalt migrated from its source in all cases of planetesimal melting. Global magma oceans and very high melt fractions almost certainly did develop within some young planetesimals. They were inferred by Taylor et al. (1993) who argued that the IVB iron meteorites crystallized from liquid iron that contained <1 wt.% sulfur (Goldstein et al. 2009) at >1770 K, a temperature at which primitive silicate material would have been well over 50% molten. Taylor et al. (1993) also noted that, if pallasite olivine represents unmelted residue, then the silicate melt fraction must have been between 70% and 90%. Keil et al. (1989) argued that the unusual texture of the Shallowater aubrite indicates a molten enstatite magma ocean at 1850 K. So while the Hevey and Sanders model is necessarily simplified, the evidence for magma oceans with high melt fractions suggests that the model is not wildly wrong. The issue of precisely how global magma oceans were created is a matter for future investigation; for now, we merely speculate that the mechanism may possibly have been linked to gradual accretion with the newly added material at an early stage (e.g., t < 1.5 Myr) continually “dissolving” in any highly radioactive rising basalt magma (see Kleine et al. 2012), or it was perhaps linked to the onset of convection before significant melting had occurred, facilitated by a possible substantial reduction in bulk viscosity (Schölling and Breuer 2009). Regardless of the details of the melting mechanism, we suspect that substantially molten interiors were the norm rather than the exception in planetesimals that accreted within field (1) of Fig. 4.