Heavy Bombardment of the Earth at -3.85 Ga: The Search for Petrographic and Geochemical Evidence

Graham Ryder Lunar and Planetary Institute

Christian Koeberl U"iversity 0/Vienna

Stephen J. Mojzsis University o/California, Los Angeles

The Moon experienced an interval of intense bombardment peaking at -3.85 ± 0.05 Ga; subsequent mare plains as old as 3.7 or 3.8 Ga are preserved. It can be assumed that the early Earth must have been subjected to an even more intense impact flux resulting from its larger size and because of its proximity to the Moon. Siderophile-element analyses (e.g., Ir abun­dance) of the oldest sediments on Earth could be used to indicate past escalated influxes of extraterrestrial material. In addition, shocked minerals may also be present in the oldest extant rocks of sedimentary origin as detrital minerals. and remnants of impact ejecta might exist in early Archean formations. Searches for impact signatures have been initiated in the oldest sedi­ments on the Earth, from the early Archean (>3.7 Ga) terrane of West Greenland; some of these rocks have been interpreted to be at least 3.8 Ga in age. So far, unequivocal evidence of a late heavy bombardment on the early Earth remains elusive. We conclude that either the sedimen­tation rate of the studied sediments was too fast and therefore too diluting to record an obvi­ous signal, or the ancient bolide flux has been overestimated, or the bombardment declined so rapidly that the Greenland sediments, some even at -3.85 Ga in age, do not overlap in time with it.

1. INTRODUCTION Earth appears to have been completed about 50-100 m.y. after the initial collapse of the solar nebula (Lee and Halli­

Collisions between planetary bodies have been funda­ day, 1995, 1996; Halliday et al.. 1996; Podosek and Ozima, mental in the evolution of the solar system. Studies under­ 2000), in a timescale apparently more protracted than that taken over the last few decades have convinced most for smaller planetesimals and Mars. As a result of later workers that the planets formed by collision and hierarchi­ geological activity, no record of any primary accretion bom­cal growth starting from small objects, i.e., from dust to bardment history remains on the surface of the Earth. planetesimals to planets (e.g., Wetherill. 1994; Taylor, Thc period on Earth between the end of accretion and I992a,b; see also chapters in section II of this volume), and the production of the oldest known crustal rocks is com­not from condensation downward. Late during the accre­ monly referred to as the Hadean Eon (Cloud, 1976. 1988; tion of the Earth (some time after -4.5 Ga), when it had Harland et aI., 1989; Taylor, 2000), which is a chronostratic reached about 70% of its eventual mass, it was most prob­ division (Fig. I). Its terminal boundary is actually not de­ably impacted by a Mars-sized or larger body (see Cameron, fined on the Earth; Harland et af. (1989) equate it with the 2000). The consequences of such an impact event for the Orientale impact on the Moon. Others do not even use the proto-Earth would have been severe and seminal, ranging teml Hadean, either distinguishing the chronometric divi­from almost complete melting and formation of a magma sions of Archean Eon and (older) Priscoan time (Harland ocean, thermal loss of preexisting atmosphere, changes in et al., 1989), provisionally at 4.0 Ga. Here we use the term spin rate and spin-axis orientation, to accretion of material Hadean to represent the time period between the formation from the impactor directly, or through rapid fall-out from of the Moon at -4.5 Ga and the beginning of the continu­orbital debris below the Roche limit. Much of the material ous terrestrial rock record at 3.8 Ga. blasted off in the impact eventually reimpacted the Earth; In contrast with the youthful age for the crust of the some of the ensuing Earth-orbiting debris would have rap­ Earth, the surface of the Moon displays abundant evidence idly coalesced to form most of the Moon and probably some of an intense bombardment at some time between its origi­smaller moonlets. Some of this geocentrically orbiting nal crustal formation and the outpourings of lava that form material would have continued to impact both bodies for the dark mare plains. Even prior to the Apollo missions, perhaps tens of millions of years after the lunar forming these plains were calculated to be about 3.6 Ga in age based impact. The essential accretion and core formation of the on crater counts and realistic flux estimates. Hence, the

475

Origin ofthe Earth and Moon Eds.: Robin Canup and Kevin Righter

University of Arizona Press, Tucson (2000)

Fig. L. Comparative chronostratigraphies of the Earth and Moon, based on Harland el al. (1989) and Wilhelms (1987). The times of interest in this paper are the Isuan and Hadean Eras for the Earth and the Pre-Nectarian, Nectarian, and Imbrian Periods for the Moon. The Imbrian is divided into the two Epochs of Early Imbrian and Late Imbrian, which have greatly differing styles of geological activity (rock stratigraphic units, i.e, systems, are not used in this paper). Although the chronostratic divisions into these two Epochs (the Nectarian and the pre-Nectarian) are perfectly clear, the cotTeiation with absolute time is less established, although the age of the Fra Mauro Formation (Imbrium ejecta morphology) that defines the division of Early Imbtian and Nectarian is fairly well established at 3.84 or 3.85 Ga (e.g., Dalrymple and Ryder, 1993).

heavy bombardment was inferred to be ancient (Hartmann, 1966). Lunar highland sample data show isotopic resetting from thermal heating, for which there is abundant evidence for impact sources dominated by ages of around 3.8-3.9 Ga. The most ancient volcanic rocks from mare plains have ages of about 3.8 Ga (see, e.g., Taylor, 1982; Wilhelms, 1987). The highland ages have been interpreted to either represent a short and intense heavy bombardment period at 3.85 ± 0.05 Ga or so (e.g., Tera et al., 1974; Ryder, 1990), or the tail end of a prolonged postaccretionaly bombardment (e.g., Baldwin, 1974; Hartmann, 1975), as discussed in Hartmann et af. (2000). In any case, the bulk of this bombardment, which produced size-seriate scars up to multiring basins many hundreds of kilometers across, preceded 3.8 Ga. We will use the term late heavy bombardment to refer specifi­cally to that bombardment of the Moon and the Earth from -3.90 to 3.80 Ga.

In any given time-span, the Earth must have been sub­jected to a significantly greater bombardment than was the Moon, as it has a larger diameter and a much larger gravi­tational cross section, thus making it an easier target to hit (e.g., Maher and Stevenson, 1988; Oberbeck and Fogleman, 1989; Zahnle and Sleep, 1997). If a late heavy bombard­ment occurred on the Moon, the Earth was subject to a flux scaling because of the ratio of the impact cross sections (Sleep et al., 1989), which may have resulted in an impact rate ~20x greater than the lunar one, containing both more and larger impact events. The consequences for the hydro­sphere, atmosphere, and even the lithosphere of Earth at that time must have been devastating (Zahnle and Sleep, 1997; Grieve, 1980; Frey. 1980). There is evidence that the Earth's upper mantle had already undergone some differentiation at the time of formation of the oldest igneous rocks, sug­gesting the prior existence of chemically evolved crust (e.g., Harper and Jacobsen, 1992; McCulloch and Bennett, 1993; Bowring and Housh, 1995). It has been suggested that the absence of any rocks older than about 3.9-4.0 Ga is the result of the ancient heavy bombardment, during which impact-induced mixing recycled early crustal fragments back into the upper mantle (e.g., Grieve, 1980; Frey, 1980; Koeberl et af.. 1998a,b). In the present contribution we outline the evidence for the character and timing of the late heavy bombardment on the Moon, and in light of this, de­scribe petrographical and geochemical attempts to investi­gate if any coeval record has been preserved on the Earth.

2. THE BOMBARDMENT HISTORY OF THE LUNAR HIGHLANDS

2.1. General

Whereas there is almost no evidence for terrestrial wit­nesses to the Hadean Eon, the pre-Nectarian Period, Nectarian Period, and the Early Imbrian Epoch cover this time interval on the Moon (e.g., Harland et af., 1989; Wilhelms, 1987) (Figs. 1 and 2). The formation of a felds­pathic crust was essentially complete by -4.44 ± 0.02 Ga,

Ryder et al.: Heavy Bombardmenl ol the Earth 477

according to the recognition oflunar ferroan anorthosite of that age. The present morphology of the highlands of the Moon reflects, almost exclusively, a history of numerous subsequent impacts that occurred prior to the extrusion of the volcanic flows that form the visible mare plains (e.g., Wilhelms, 1984, 1987). These ancient impact structures in­clude giant multiring basins and their debris (Spudis, 1993), as well as a size-seriate range of smaller craters. Hartmann (1965, 1966) recognized that most ofthis cratering occurred early in lunar history according to an estimate of the aver­age age of mare plains of 3.6 Ga, which was calculated based on present-day cratering rates. He inferred a cratering rate averaging roughly 200x higher for the first one-seventh of lunar history than for the remainder. The general correct­ness of Hartmann's conclusion was demonstrated by the return ofApollo samples, and the dating of the oldest mare plains at close to 3.8 Ga (Wilhelms, 1987).

Geochronological studies of impact-brecciated highland samples show thermal events, most of them of impact ori­gin, concentrated at -3.8-3.9 Ga. These ages have been taken to represent the tail end of a heavy but declining bombardment dating back to the accretion of the Moon (e.g., Shoemaker, 1972, 1977; Hartmann, 1975. 1980; Neukum et af., 1975; Baldwin, 1971, 1974. 1981, 1987; Taylor, 1982; Wilhelms, 1987); alternatively, they may record a sharp or cataclysmic increase in bombardment for that short interval (e.g., Tera et aI., 1974; Ryde/; 1990; Dalrymple and Ryder, 1993, 1996). There exists a sharp drop-off in estimates for the cratering rate from the youngest high­land surfaces, the Orientale and Imbrium ejecta blankets, to the oldest mare surfaces. This is according to crater counts of those surfaces, which differ by a factor of -3-4 (e.g., Wetherill. 1977. 1981; B VSp, 1981). As a result of the difference in cratering rates, a flux at least lOOx higher can be calculated for this transition period, even if those young­est highland surfaces are as much as 100 m.y. older than the oldest mare plains, which have been collecting craters for 3.8 G.y. With new studies that have expanded the age ranges for the oldest known rocks on Earth, the time span for lunar bombardment now overlaps that of these oldest rocks. Therefore, a more detailed look at the chronology and intensity of the lunar bombardment can help to understand the conditions on Earth at the time of life's emergence (Mojzsis et aI., 1999). The reader is referred to the paper by Hartmann et al. (2000) for a more complete discussion of lunar cratering history.

2.2. Relative and Absolute Ages of Highland Stratigraphy

The stratigraphy of the lunar highlands has been divided on the basis of basin fonnation and ejecta into pre-Nectarian Period, Nectarian Period, and Early Imbrian Epoch (Wilhelms, 1984, 1987) (Figs. I and 2). These are separated by the time of production of the Nectaris Basin deposits, the Imbrium Basin deposits, and the debris blanket of the Orientale Basin respectively. Several basins were produced

478 Origin of the Earth and Moon

Magma ocean

Oldest ferroan anorthosites/ crustal formation

Apollo 11 mare lavas

Apollo 17 mare lavas

** = age of oldest Akilia sediment

.........~ Numerous basins South Pole-Aitken basin

. 1# jCProcellarum basin?

/ /'A-P-O-II-0-1-5-A-p-e-n-n-in-e--'V

/ / Bench KREEP basalts ?? /; . { (3.84)

/ ~rustal '-m-a-g-m-a-t-is-m--------'

/ and volcanism

?~ 1/.

Late Imbrian

3.7

!III!!l~!III!!lt:::::".~orientale basin (3.80/3.84?) =: Schr6dinger basin

~~~.-.iL~ -"4__.lmbrium basin (3.85) " Serenitatis basin (3.89)

3.9 -+-----------------------'---1--.. Crisium basin (3.89) Nectaris basin (3.90/3.92?)

.~-??--....-

Ga 3.6 -,---------r­

3.8

4.3

4.5

4.0

4.2

4.1

4.4

Fig. 2. Stratigraphy and chronology of early lunar history, based on relative stratigraphy discussed in Wilhelms (1987) and absolute agc inferences as discussed in this paper. The basins with underlined names define the stratigraphic column. Some other significant events or features of early lunar history are shown. While significant impacting and contraction of the geologic column is obvious at 3.8-3.9 Ga, the event/time correlations within the pre-Nectarian, and even the age of the Nectaris Basin, are much more contentious. The age of the oldest Akilia sediments, discussed in this paper, are shown (**) for comparison with lunar stratigraphy.

during the Nectarian Period, including Serenitatis and Cri­sium, whose ejecta regions have been sampled (Apollo 17 and Luna 20). The Schr6dinger Basin is Early Imbrian, as are several large craters, including some that are almost 200 Ian in diameter. The oldest mare deposits were erupted in the Late Imbrian Epoch, the end of which is defined in terms of crater degradation and crater counts in the absence of any globally useful stratigraphic-datum horizons compa­rable to basin ejecta. The dating of these boundaries, as well as of other basins within the stratigraphic units, defines the chronology of lunar bombardment and the flux over the main period of interest here. Absolute ages quoted have been recalculated using the revised decay constants of Steiger and

Jager (1977), and thus most are slightly younger than those given in some of the original publications.

Although these divisions for lunar time were introduced above in normal stratigraphic sequence from oldest to youngest, it is more convenient to discuss the absolute dat­ing of the boundaries from youngest to oldest, from the sim­plest interpretations based on the best preserved impact craters, to the more difficult.

2.2.1. The oldest mare surfaces. The Late Imbrian Epoch commenced with the formation of the Orientale Basin, the final large multiring basin to have formed on the Moon. It was followed by few large (> IO kIn) cratering events. The end of the Late Imbrian Epoch is arbitrarily

defined, and includes the mare basalts at the Apollo IS land­ing site that have been dated at -3.25 Ga. Lavas dating to the Late Imbrian compose roughly two-thirds of the mare surfaces. Most important for the discussion here are the older mare units, including those from which mare basalt samples were collected at the Apollo II and Apollo 17 land­ing sites. The common Apollo II group B2 mare basalts and the rare ApolIo II group D mare basalts are 3.80 Ga or slightly older; some Ar-Ar age determinations are as old as 3.85 Ga (Snyder et al., 1994, 1996). At the Apollo 17 land­ing site, the oldest mare basalt so far identified formed at 3.87 ± 0.10 Ga (Dasch et al., 1998), and other mare basalts from there are almost 3.80 G.y. old (see summary in Wil­helms, 1987). Although younger basalts were also collected from these locations, it seems likely that the presence of such old basalts close enough to the surface to be in the sample collection suggests that for all but the smallest cra­ters (those that are a few meters across) the crater counts for these areas represent surfaces very close to 3.80 G.y. old, and perhaps slightly older. At a minimum, the crater counts for these sites represent surfaces that are at least 3.6, and probably more than 3.7, G.y. old.

2.2.2. The age of the beginning of the Late Imbrian Epoch (the age of Orientale). The Orientale and Schri:i­dinger Basins are far removed from any sites sampled so far. Their ejecta have crater counts that are similar to each other and they are slightly less cratered than the Imbrium ejecta (Fra Mauro Formation, Cayley Formation). Schri:i­dinger is older than Orientale, as it is superposed by Orientale secondaries. However, their absolute ages cannot (yet) be independently dated; they are older than the oldest affected mare plains, and thus are construed as older than -3.80 Ga.

2.2.3. The age of the beginning of the Early Imbrian Epoch (the age ofImbrium). The best way to date an im­pact is by using the radiogenic isotopes in a clast-free or clast-poor impact-melt rock (Ryde/; 1990; Deutsch and Scharer, 1994). Unfortunately, impact-melt rock that can be identified specificalIy as a product of the Imbrium impact is lacking, and those considered most likely (the Apollo 15 dimict breccias; Ryder and Bower, 1977) have disturbed Ar­Ar systems (e.g., Bogard et aI., 1991). Recently, Haskin et al. (1998) have argued that all Th-rich impact melt brec­cias (low-K Fra Mauro) collected 011 the Apollo missions are products of the Imbrium impact event. Despite the dat­ing problems, there are ways to bracket the age of Imbrium. First, the Apennine Bench Formation is a volcanic plateau inside (hence younger than) the Imbrium Basin. Remote sensing of its morphology and chemistry allow correlation with Apollo IS volcanic KREEP basalt samples, which have been dated. This provides a lower age limit on the Imbrium Basin of 3.84 ± 0.02 Ga (Ryder. 1994). Second, the Apennine Front has been little modified since the forma­tion of the Imbrium Basin. Thus crystalline impact melt collected there should be almost entirely Imbrium impact melt, or older impact melt. Dalrvmple and Ryder (1991, 1993) dated such melt rocks and suggested that Imbrium is

Ryder et al.: Heavy Bombardment of the Earth 479

certainly no older than 3.870 ± 0.010 Ga, and probably no older than 3.836 ± 0.016 Ga. Third, similar arguments ap­plied to the contents of the Fra Mauro Formation and the Cayley Formation suggest a similar age constraint. For ex­ample, impact-melt fragments in the white rock 14063 from Cone Crater show a range from 3.87 to 3.95 Ga; other samples that are probably not from the Fra Mauro Forma­tion, but represent later local events (such as the 14310­group samples), are a little younger (3.82 ± 0.02 Ga). Thus, it is safe to bracket Imbrium as 3.85 ± 0.02 Ga. This is consistent with the the older Serenitatis Basin (below) hav­ing formed at about 3.89 Ga.

2.2.4. The age ofthe beginning of the Nectarian Period (the age of Nectaris). Stratigraphic and crater count data show that the Nectaris Basin is older than the Crisium Ba­sin, but melt-rock samples from it cannot be identified with certainty. The Apollo 16 site was modified by Nectaris ejecta, and subsequently by Imbrium ejecta. Fragments within the breccias collected on the Apollo 16 mission prob­ably include samples of melt created prior to Nectaris in several clearly recognizable large craters that underlie the site. None of the melt samples dated so far is reliably older than 3.92 Ga. The analysis of the rocks and ages by James (1981) strongly suggests an age for Nectaris of less than 3.92 Ga, and probably an age of -3.90 Ga is reasonable, consistent with earlier derivations by Turner and Cadogan (1975) and Maurer et al. (1978).

The Nectalian Period also witnessed the formation of the Serenitatis and Crisium Basins, and samples were collected from their rims or ejecta. At the Apollo 17 landing site, the highland materials are dominated by coherent poikilitic melt rock, commonly in the form of boulders whose trails can be seen to run high up the massifs. These samples are most readily interpreted as melt formed in the Serenitatis Basin event. If they are not, then they are probably older, as it is inconceivable that they are balIistic ejecta from the Imbrium event. Most of these samples belong to one chemical group whose age, as determined on several samples, is now pre­cisely established as 3.893 ± 0.009 Ga (Dalrymple and Ryder, 1996). This age is outside of the bracket for the Imbrium age described in the previous section. The Luna 20 sample from Crisium ejecta includes impact-melt rock samples. From these, Swindle et al. (1991) suggested an age of -3.89 Ga for the Crisium Basin. These ages for Serenitatis and Crisium are consistent with an age of the older Nectaris Basin of3.90 Ga. Several other basins, e.g., Herzsprung and Humorum, also formed after Nectaris. Thus, there was con­siderable bombardment of the Moon in the 60 m.y. between 3.90 Ga and -3.84 Ga.

2.2.5. Pre-Nectarian Period and events. The lack of impact-melt rocks in the sample collections that are older than -3.92 Ga cannot be due to resetting of all older ages, given the difficulties of such resetting (e.g., Ryder, 1990; Deutsch and Scharer, 1994). Most of the lunar upper crust has not been converted into impact-melt rock, which would be subject to resetting. Thus the paucity of pre-3.92-Ga impact melt can be taken as evidence that there was little

480 Origin of the Earth and Moon

impacting prior to that time, other than that expressed by the metamorphosed breccias of uncertain origin, the felds­pathic granulites, that may well date back to the very earli­est postaccretionary bombardment at about 4.4 Ga. Further­more, the Pb-isotopic data of Tera et al. (1974) indicate events at -3.85 Ga and events at >4.4 Ga, but not much evidence of events in between; continual resetting of Pb clocks would show up as intersects in the 4.4-3.9-Ga Pb­isotope growth curve. There is also a lack of the comple­ment of siderophile elements that would be expected to be present in older upper crustal rocks if a heavy bombardment between 4.4 and 3.9 Ga had occurred (Ryder, 1999), despite claims to the contrary (e.g., Sleep et al., 1989; Chyba, 1991). A more complete discussion of these features appears in the chapter by Hartmann et at. (2000).

2.3. Summary of the Significance of the Lunar Cratering Record from 3.90 to 3.80 Ga

During the period from 3.90 to 3.80 Ga, a substantial amount of the extant lunar highland features, including Nectaris and many younger basins, formed on the Moon. Based on the above discussion, it is possible that this in­tense activity terminated at 3.85 Ga with the near-simulta­neous creation of the Imbrium, Schr6dinger, and Orientale Basins. Bombardment may even have finished as early as 3.87 Ga. Although the last two basins might have fOlmed as late as 3.80 Ga, this seems unlikely given that their su­perposed crater populations are almost as high as those on the Imbrium ejecta, yet greater than those on the oldest mare plains, which themselves are -3.8 G.y. old. Many lunar basins, including South Pole-Aitken, formed prior to this period, but at present there is no direct or definitive way to date their formation; South Pole-Aitken might be as young as 3.95 Ga, or as old as 4.3 Ga. In principle, future missions can obtain samples from which the chronology of ancient intense bombardment can be more reliably determined. In particular, the age of Orientale could be precisely deter­mined, as its impact melt sheet is intact and accessible, and would constrain the younger end of bombardment. South Pole-Aitken may be datable and could provide a constraint at stratigraphically older times, because although it has been battered, remnants of the impact-melt sheet should be col­lectable and recognizable.

The chronology outlined above suggests a massive de­cline in the flux of bombardment on the Moon over a short period of time following the Nectaris event. The cratering on the Nectaris ejecta (3.90 Ga) is a factor of -4 higher than that on Imbrium ejecta (3.85 Ga), which, in tum, is a fac­tor of -4x that recorded on the oldest mare plains (about 3.80 Ga, or even slightly older).

As an exercise, let us define that there are C units of craters on 3.80-G.y.-old mare plains. Then the average cratering rate from 3.80 Ga to the present is C per 3.80 units and Ga (i.e., 0.263C units/Ga). Imbrium ejecta has -4C units of craters. Thus, the cratering rate between Imbrium forma­tion and oldest mare plains is (4C-IC) per 0.05 units and Ga (assuming 50-m.y. age differences), which is a rate of

60C units/Ga. Therefore, the relative cratering rate of the 3.85-3.80-Ga period compared with the average since 3.80 Ga is -228.

Furthermore, let us assume that Nectaris is 3.90 Ga and Imbrium is 3.85 Ga. There are -16C units of craters on Nec­taris ejecta. Thus the cratering rate during this period is (l6C-4C) per 0.05 units and Ga, i.e., 240 units/Ga. This is 912x the average rate since 3.80 Ga.

The present rate, or the Phanerozoic rate, of cratering is probably a little lower than the average over the last 3.80 G.y., because there was a higher flux in the Late Imbrian Epoch than in the succeeding Eratosthenian and Copernican (in­deed, there is evidence that suggests a higher flux in the Eratosthenian than in the Copernican; Ryder et al., 1991; Culler et al., 1999). In round figures the cratering rate in the period 3.90 Ga to 3.85 Ga was probably at least 1000­1500x that of the Phanerozoic, and in the period 3.85-3.80 Ga was probably at least 250-400x that of the Phanero­zoic. These are higher than the rates inferred by Hartmann (1966), because he assumed that the observed cratering record stretched back almost to the origin of the Moon, whereas it is actually much more restricted in time. It is even possible that the later decline took place over only the first 10 or 20 m.y. after -3.85 Ga, such that by 3.84 Ga or 3.83 Ga the flux was approaching within a few factors of that of the present day. This is the record that needs to be compared with that of the oldest rocks on Earth.

3. STATE OF THE SURFACE OF THE EARTH FROM 4.5-3.8 GA

3.1. Earliest Crust

Recognized extant terrestrial crustal rocks extend back to only about 89% of the history of the planet, to -4.0 Ga; the record is improved somewhat if we include the poten­tial information gleaned from rare detrital zircon grains that are up to 4.27 Ga in age, -94% of Earth history. As men­tioned above, is likely that the Moon-forming impact led to a large-scale melting of the Earth and the existence of an early magma ocean (e.g., papers in this volume). Mantle temperatures in the Hadean were probably much higher than today. About half of all heat produced by 235U decay to z07Pb was released during the Hadean, 40K was more abun­dant, as well as latent heat from accretion, all of which added several hundred degrees to the internal temperature of the Earth. Any late accretionary bodies would have added further thermal energy to the already elevated budget of heat flow in the early Earth (e.g., Smith. 1981; Davies, 1985; Taylm; 1993. 2000).

The nature of the earliest crust on Earth, and the amount of crust present, has been the subject of intense debate. Petrological melting concepts and comparisons with other planets suggests that the earliest crust on Earth was basal­tic in composition (e.g., Taylor, 1989. 1992. 1993; Arndt and Chauvel. 1991). The existence of any substantial early feld­spathic crust on the Earth seems precluded by the higher pressure at shallower depth on the Earth (in contrast to the

Moon, which does have an ancient feldspathic crust). Cal­cium and Al are sequestered in the deeper Earth in early high-pressure phases (particularly garnet), which delays the concentration of those elements reaching the point of pla­gioclase crystallization. In addition, plagioclase itself can­not crystallize at significant pressure (hence depth), so any plagioclase-bearing terrestrial crust would be thin. Neither would any plagioclase be likely to float in a water-bearing, basaltic magma ocean on the Earth, so no concentration of plagioclase toward the surface would be realized. Finally, there is no indication of any ancient reservoir of Eu or of primitive 87Sr/86Sr signatures that could have resided in an early high-Sr and low-Rb anorthositic crust (e.g., Taylor, 1989). Some pre-4.0-Ga differentiation of the mantle seems to have occurred, as indicated by isotopic evidence (e.g., Hmper and Jacobsen, 1992; Bowring and Housh, 1995; but see also Gruau et aI., 1996, for cautionary remarks). Detrital zircon crystals in an Archean quartz-pebble conglomerate from the Narryer Gneiss Complex, Western Australia, are the oldest known minerals on Earth, with ages up to 4.27 Ga (e.g., Compston and Pidgeon. 1986). The morphological, mineralogical, and geochemical characteristics, as well as similarities with post-3.75-Ga zircons, indicate a compos­ite granitoid source of continental provenance for these zir­cons (Maas et al., 1992; Mojzsis, 1998). Thus there is evi­dence for at least minor amounts of felsic igneous rocks in the Hadean, which may have been present in small amounts from remelting of basaltic crust that sank back into the mantle (e.g., Taylor, 1989, 2000). lt remains unlikely, though unproven, that significant amounts of continental crust ex­isted on Earth during much of the Hadean Eon. The lack of initial Hf-isotopic heterogeneity and the absence of nega­tive cHf values in early Archean rocks provides evidence against the presence of large amounts of continental crust on the Hadean Earth (Vervoort and BUchert-Tojt, 1999). Gra­nitic crusts require multistep derivation from the primitive mantle by recycling of subducted basaltic crust through a "wet" mantle, which will slowly lead to an increasing amount of granitic crust through time (Taylor and McLennan, 1995).

The lithosphere of the Hadean Earth was most probably characterized by a basaltic crust, covered by an ocean, and with little dry land and only minor amounts of felsic rocks (granitoids). Any sedimentological record, which would host infOimation specific to surface environments such as the rate and violence of meteorite impact and the presence of life, has been almost completely lost from Hadean times, and only appears at its conclusion, near 3.90 Ga (Mojzsis et al., 1996. 1999; Mojzsis and Harrison, 2000; Nutman et al., 1996, I997).

3.2. Effects of Ancient Impacting: From Basins to Dust

Individual impacts have considerable physical (morpho­logical) and chemical effects on the target and on the at­mosphere. A crater is excavated, fragmental ejecta are strewn around and into the crater, and a melt unit can be created. Some of the ejecta might be in the form of molten

Ryder et a/.: Heavy Bombardmenr of/he Earth 481

spherules. The projectile and some target rock are vapor­ized, and a fraction of the projectile vapor can be incorpo­rated into melt and ejecta. Minerals of both the autoch­thonous target and the allochthonous ejecta could exhibit shock effects (e.g., planar deformation features, high-pres­sure polymorphs, diaplectic glasses) from the interaction of the rocks and minerals with the shock wave. If the target includes water (e.g., ocean impact) then that water gets vaporized. If the impactor contains typical chondrite-like abundances of platinum-group metals (e.g., Ir), as all me­teorites other than most differentiated stony meteorites do, then these will be added to the impactite. Depending on the density of the atmosphere, there is a lower size limit below which small impactors do not penetrate the atmosphere and therefore will not form craters. At very large impactor di­ameters, excavation of mantle material is possible, as well as large-scale vaporization (and possible loss) of atmosphere and hydrosphere. Judging from the lunar record (see sec­tion 2), very little of the Earth's surface in the period of -3.9-3.8 Ga should have escaped being the target of sig­nificant impacts at one time or another, and therefore es­caped being covered by ejecta from craters that are at least a few kilometers in diameter. However, a more vigorous rock cycle than at present continually resurfaced the early Earth and erased (most?) evidence for such an impact en­vironment.

Calculations that scale impact-melt production with in­creasing crater dimension (Melosh, 1989; Cintala and Grieve. 1994, 1998) show a breakdown of this geometric relation­ship for very large impact structures. As the magnitude of the impact increases, the melt volume relative to the tran­sient crater size increases, with a larger proportion being retained inside the crater, and the depth of melting for large impact structures exceeding the depth of excavation. There­fore, the thermal effects of an impact (i.e., the large-scale melting) will actually reduce the amount of shocked rocks that are formed and preserved. In large-scale impact events, leading to the formation of craters larger than a few hun­dred kilometers in diameter, thermal metamorphism may be more important than shock metamorphism. However, cra­ters smaller than a few hundred kilometers in diameter would still largely have fragmental and shocked ejecta and basement.

Most of the (considerable) speculation regarding the ef­fects of ancient impacts on the Earth has focused on large, potentially basin-forming, events. These models attempt to understand the localization and extent of endogenic activ­ity, such as volcanism, proto-ocean basin formation, atmo­spheric disturbance, continental growth and assembly, and changes in sedimentation style and topography, rather than relying on direct impact evidence. Grieve (1980) and Frey (1980) discussed the effect of impact structures with diam­eters exceeding 100 km on the ancient Earth, prior to about 3.8 Ga. By scaling the lunar impact record to the Earth these authors concluded that about 2500-3000 impact structures with diameters larger than about 100 km could have formed. Their simulation resulted in almost 1000 craters with diam­eters exceeding 200 km, and possibly about 10 structures

482 Origin of the Earth and Moon

with diameters larger than that of the Imbrium Basin on the Moon (about 1300 km diameter). This crater population would have covered about 40% of the surface of the Earth. Using the minimum estimate for the cratering frequency, Grieve (1980) derived a cumulative energy of about 1029 J added to the Hadean Earth from impact events, and con­cluded that the net effect of large impact events was to localize and accelerate a variety of endogenic geological activity.

Several studies have considered the effects of impact on the atmosphere and hydrosphere, again, particularly for very large events (Maher and Stevenson, 1988; Oberbeck and Fogleman, 1989; Sleep et al., 1989; Chyba, 1993; Zahnle and Sleep, 1997). These studies have largely been expressed in the context of the early evolution of life and impact-in­duced sterilization. An Imbrium-scale impact onto the early Earth would have the ultimate effect of boiling off about 40 m of seawater, with a subsequent hot surface layer and annihilation of any surface ecosystems (Zahnle and Sleep, 1997); expected events lOx as large as this would have correspondingly larger and more devastating effects. It prob­ably requires the impact of an asteroid several hundred ki­lometers in diameter to totally vaporize one present-day ocean mass of water. The scale of these events is probably too great and destructive to allow preservation of evidence. It is the probability for these vaporizing impacts on the early Earth that has led to the general impression that impact events were a negative forcing function for the development and evolution of emergent life (e.g., Grieve, 1998).

Along with the mega-impacts there would be numerous smaller impacts, producing more recognizable ejecta blan­kets, shock features, and input of siderophile elements. Si­multaneously, there should be a correspondingly greater abundance of input of interplanetary particles and continu­ous rain of dust (that ultimately is incorporated into rocks with ongoing sedimentation) than there is at the present day. It is to these smaller-scale features that attention should be paid, to find evidence of impact in the oldest rocks.

4. SEARCH FOR EVIDENCE OF A LATE HEAVY BOMBARDMENT ON

THE EARLY EARTH

4.1. Earliest Sedimentary Rocks on Earth

The critical sedimentary record of the earliest Archean is preserved in the North Atlantic province, principally in the Isua district and the Akilia association in southem West Greenland that are part of the Itsaq Gneiss complex (Nutman et aI., 1996). The Isua Supracrustal belt is in effect a giant version of the smaller enclaves of Akilia rocks with abun­dant gneisses. The Itsaq Gneiss complex of West Greenland is a 3000-krn2 terrane dominated by orthogneisses of grani­toid compositions that intrude, in some locations. packages of associated sediments and volcanic rocks. These supra­crustal rocks are composed of massive amphibolites and complex metasomatic carbonates (metamorphosed equiva­lents of pillow basalts and other components of early

Archean oceanic crust), ubiquitous banded iron formations (chemical sedimentary precipitate, dominated by quartz and magnetite), rare graywacke, and metapelites. The inferred environment of deposition for these volcanosedimentary successions is a sediment-poor arc or back-arc basin in rela­tively deep water (Nutman et aI., 1984). Studies of early Archean sediments from the Isua Supracrustal belt (ISB), which are -3.80 Ga, and rocks of the Akilia association in the Godthabsfjord region (>3.80 Ga) of southern West Greenland, suggested that they are the oldest sediments yet identified (Nutman et al., 1997).

There are uncertainties concerning geological relation­ships on Akilia island and the nearby islets of the Godthabs­fjord archipelago that host the oldest known sediments of marine origin, and also contain evidence for life (Mojzsis et al., 1996; Natman et aI., 1996, 1997). These derive from reconnaissance-scale geological mapping that reveals little about the structural relationship of the banded iron forma­tions to the polyphase, geochemically heterogeneous ortho­gneisses that intrude them. The geochronological relation­ships as they are currently inferred have been used to place a minimum age of formation for some of the sediments in excess of 3.85 Ga (Nutman et aI., 1997). In contrast, Moor­bath and co-workers (e.g., Moorbath et aI., 1997; Kamber et aI., 1998; Moorbath and Kamber, 1998; Kamber and Moorbath, 1998; see also Rosing, 1999) argued that these (;

oldest ages represent only those of zircon inherited from as­similated preexisting rocks older than the intruding grani­toid orthogneisses on Akilia and the surrounding islands. However, evidence for much Pb contamination from hypo­thetical assimilated zirconiferous rocks is absent from the orthogneisses of the Itsaq, so the zircons are probably not inherited. Furthermore, the intruding granitoids are low in Zr, granodioritic melts are strongly undersaturated with re­spect to Zr, and the rocks they intrude are poor in zircon. Age estimates of 3.65 Ga derived for the intruding gneisses of southern West Greenland, which are based on whole-rock Pb-Pb, Sm-Nd, and Rb-Sr errorchrons, are susceptible to open-system REE-, Sr-, and Pb-diffusion behavior, in con­trast to precise and concordant zircon geochronology (Mojz­sis and Harrison, 2000). It is not possible to resolve the age issue here; however, this question has important implications for the search of traces of any late heavy bombardment on the Earth, as these terranes presently provide the only quali­fied samples to search for extraterrestrial components of a late heavy bombardment on Earth. We recognize that the evidence for a 3.85-Ga or older age for the sedimentary Akilia rocks under consideration here is stronger than that for a younger age.

4.2. Search Strategies and Their Rationales

We discuss three strategies used to search for evidence of a late heavy bombardment on the early Earth. First, it is possible to search for chemical evidence in sedimentary rocks that would indicate an enhanced flux of extraterres­trial materials, using different techniques and samples from both Isua and Akilia rocks from Greenland. Second, evi­

dence of detrital shocked minerals that might have formed as a result of an inccssant early bombardment may be pre­served. Third, it may be possible to recognize remnants of impact ejecta (albeit strongly altered and metamorphosed) that might have been incorporated into early Archean rock formations.

4.3. Meteoritic Siderophile-Element Signatures

4.3.1. Siderophile elements on the early Earth. The Earth is a highly differentiated body, with a core, a mantle, and evolved crust. During planet formation, the highly sidero­phile elements (e.g., Ir, Pt, Au) partition strongly into me­tallic cores. The formation of the Earth's core was completed early, well before the formation of the most ancient of pre­servcd terrestrial crustal rocks, and certainly by the time of lunar formation. Thus, Earth's earliest mantle and crustal rocks were effectively stripped of their highly siderophile' element inventory early on. However, the present-day up­per mantle has abundances of highly siderophile elements much higher than expected from presently known silicate­metal distribution coefficients and under the assumption of core-mantle equilibrium (lr -3 ppb) (Chou, 1978; Chou et al., 1983; Newsom, 1990). The siderophile-element abun­dances show chondritic relative proportions, which plot subparallel to thc Clline. The addition of-0.75% chondritic material after tennination of the core-upper mantle equilib­rium under increasingly oxidizing upper mantle conditions seems necessary to explain the abundances and chondritic relative proportions of the siderophile elements in the mantle (e.g., Chou et al., 1983; Newsom, 1990; Holzheid and Palme, 1998). The emplacement timing of such a veneer is not constrained by direct evidence, but is often invoked to have been as early as 4.40 Ga, or as late as 3.80 Ga.

Siderophile elements are strongly fractionated during partial melting; for example, basalts are strongly depleted in Ir «0.05 ppb) relative to mantle peridotites (-3 ppb). In rare circumstances, siderophile elements can be concen­trated in specific crustal reservoirs, e.g., platinum-rich lay­ers in some basic intrusions; these have relative platinum­group element abundances that are strongly fractionated from chondri tic values. More evolved rocks, such as pyroclastics, granites, and the sediments derived from them, contain negligible siderophile-element abundances from terrestrial sources. The source of significant abundances of siderophile elements in evolved crustal rocks, such as the melt rock at East Clearwater Lake crater (Palme et al., 1979), or in the Cretaceous-Tertiary boundary clay layer (Alvarez et al., 1980), can be reliably attributed to an ex­traterrestrial source. Thus siderophile elements in sedimen­tary rocks, other than in the rarest of circumstances, can be taken as an indication of an extraterrestrial flux at the time of formation of the sediments, particularly if they are in chondritic relative abundances.

Estimates of the flux of extraterrestrial material to the Earth, based on lunar stratigraphic-chronologic studies dis­cussed above, suggest that during the peak of the late heavy bombardment this flux was, or ranged, between -3 x 102

Ryder et al.: Heavy Bombardment of the Earth 483

(low estimate) and -104 (high estimate) greater than at present. While this estimate is based on visible lunar cra­ters, generally of the order of a few kilometers in diameter and larger, it is inferably true of smaller craters and of in­terplanetary particles and dust as welL In a geochemical sense, it does not matter whether a projectile makes a cra­ter or bums up in the atmosphere; it wil\ be added to the sediment as the dust settles. On the Moon the extralunar material also has high abundances of the siderophile ele­ments (for example, the Serenitatis impactor was almost certainly an EH chondrite, James, 1995). All lunar impact­melt rocks from the late heavy bombardment contain Ir in the 2-20-ppb range (Papike et al., 1998). On Earth, a sedi­mentary layer at -3.85 Ga might show evidence for an in­flux of siderophile elements from an enhanced continuous background fallout, or from a specific event comparable with the Cretaceous-Tertiary boundary layer, where such events had a higher probability than at the present.

4.3.2. Terrestrial sources ofiridium in marine sediments. Experiments have shown that -50% of Ir in sediments is scavenged from seawater by Fe-Mn-O-OH particles (Anbar et aI., 1996) in oxic to suboxic environments. Anoxic envi­ronments, such as would be the case for much of the Archean hydrosphere (Holland, 1984), are not a major sink for Ir because of the redissolution of particulate hydroxides, except at rapid sedimentation and relatively shallow water depths. Iridium is well mixed in the oceans: The residence time for it in the hydrosphere is 2000-20,000 yr. This im­plies that extraterrestriallr could persist in seawater and be incorporated into sediments by particulate scavenging be­tween impacts of a frequency of Jess than 2000 yr. It was also found by Anbar et al. (1996) that Ir (and as) abun­dances in present-day seawater are extremely low. Thus weathering and hydrothermal alteration of ultramafic rocks, such as peridotite, which could supply Ir (and other plati­num-group elements) to seawater, is insignificant in deter­mining the abundances of these elements in present-day seawater.

Studies of REE distributions in banded iron formations demonstrate that hydrothermal activity had a strong influ­ence on overall seawater chemistry in the Archean (Bau and Moller, 1993). The average concentrations of lr in pelagic clays with sedimentation rates of -0.001-0.003 mm a-I range from 0.07 to 2.0 ppb (Barker and Anders, 1968; Kyte and Wasson, 1986); in metalliferous sediments that scavenge Ir, concentrations are even higher (Anbar et aI., 1996). Some of these higher abundances might result from organic mat­ter scavenging, and therefore do not reflect extraterrestrial input directly. Because these pelagic sediments are very slow to accumulate, they contain measurable lr even at the present-day very low rates of meteoritic input.

4.3.3. Extraterrestrial sources or iridium to the hydro­sphere. Estimates of the influx of extraterrestrial matter reaching Earth's surface during the past 100 m.y. have been the subject of numerous studies aimed at quantifying the current rate of dust accretion and the composition and source of the materiaL A number of methods have been used to determine this flux, using the collection of dust in the

484 Origin of the Earth and Moon

atmosphere, glacial ice, and pelagic sediments (Love and Brownlee, 1993). The current mass flux of extraterrestrial Ir is based on measurements of sedimentary Ir in systems with calculable sedimentation rates and calculations of the flux of infalling dusts by satellite, radar, and airborne ob­servations. Love and Brownlee (1993) have estimated the amount of chondritic material raining into the Earth as dust, from measurements of abundances and sizes of microcraters developed on the Long Duration Exposure Facility (LDEF) experiment, as 40 (± 20) x 109 g a-I. Assuming chondritic relative proportions, this translates to 70 ± 35 mol Ir a-I to the whole Earth. The uncertainties reflect counting and, more importantly, the inferred encounter velocities. An as­sumption is that the six-year length of the LDEF experiment is adequate to be representative of the current (i.e., last few million years) flux and its possible variations. The abun­dances are consistent with those derived from Os isotopes in deep-sea sediments (Esser and Turekian, 1988) and lr in both Antarctic ice (Ganapathy, 1983) and abyssal red clays (Kyte and Wasson, 1986). Studies by Bonte et al. (1987) have shown that almost all platinum-group elements present as cosmic debris occur in grains «I0 ~m. These grains are quickly incorporated into sediments (Esser and Turekian, 1988). The finer dust grains are probably sensitive to sea­water oxidation and hydrolysis after burial and have prob­ably always contributed to a small hydrogenous component of seawater Ir.

4.3.4. Ancient sediments and model extraterrestrial influx. The oldest terrestrial sediments might be expected to pre­serve a signal of higher incident fluxes from interplanetary dust particles, micrometeorites, local impacts, airburst, cometary showers, and ablation products of such phenom­ena. The amount of Ir from the background that would be expected to be sampled by the water column and thus a sediment deposited or precipitated from the early Archean ocean, [Ir]SED' can be estimated by

where cl>m = estimated present extraterrestrial flux for all incoming material, f = factor increase for ancient flux, [Ir]ET = concentration oflr in extraterrestrial material, ffiA= Earth surface area, <JlSED = sedimentation rate of the deposit SED, and PSED = density of SED.

A critical parameter in equation (I), other than those already discussed, is the sedimentation rate for the sample being analyzed for siderophile elements. Different sediments have deposition rates that differ by orders of magnitude. For our purposes, samples with a slow deposition rate are de­sirable, because they have a higher proportion of extrater­restrial material (which is why clay was investigated at the Cretaceous-Tertiary boundary; Alvarez et aI., 1980). The early Archean sedimentary rocks we have to work with are autochthonous precipitates, such as banded iron formations and quartzites (as metamorphosed chert). They contain neg­ligible contributions from the weathering detritus of igne­ous rocks (Dymek and Klein, 1988), which is advantageous insofar as some of these (e.g., peridotites) could blur the

siderophile-element signal from hydrogenous sources of Ir (Anbar et aI., 1996) and other metals, although this is not a major concern. Banded iron formations from the Isua dis­trict of southern West Greenland, and also younger ones from West Australia and southern Africa, contain no signifi­cant clastic sediment components and no near-shore or evaporitic facies. Therefore, when the oldest banded iron fornlations formed, they must have sampled for the most part the chemistry of the water column from which they precipitated, including any extraterrestrial component.

Unfortunately, the sedimentation rate for banded-iron­formation deposition is poorly constrained. These rocks do not form in Phanerozoic environments because the p02 of the atmosphere has been too high since the Proterozoic Era, and because Fe2+forming by rapid oxidation to Fe2+(Fe3+h 0(OH)6 transforms to Fe2+(Fe3+)204 (magnetite), which has low solubility in seawater. In general, detailed sedimento­logical interpretations of banded-iron-formation sequences have not been available (Klein and Beukes, 1990). There have been considerable differences of opinion about the origin of banded iron formations and the particular envi­ronments of their deposition (James, 1954; Trendall and Blackley, 1970; Cloud, 1973; Holland, 1973; Klein and Beukes, 1990), with general agreement that deposition took place below wave-base. The individual bands of iron for­mations are considered by many workers as being equiva­lent to varves associated with seasonal changes in upwelling, productivity, and local 0 production and other factors (Hol­land, 1984, and references therein). Trendall and Blackley (1970) estimated the rate of deposition of the Hamersley banded iron formation (-2.5 Ga). From counting chert + magnetite ± hematite microband couplets between volcanic rocks of known age that were interbedded with the iron­stones, these authors estimated a deposition rate of 0.65­1.3 mm a-I, which is much faster than even typical detrital sediments such as shale and siltstone. However, while band­ing in the Hamersley Basin may be on a scale of - I mm, banding elsewhere is on much coarser (centimeters) and much finer (submillimeter) scales. Indeed, banding occurs at various repetitions in any sequence, with laminae bundled into alternatively quartz-rich and magnetite-rich "beds" and higher-order packages. Ifbanded iron formations are domi­nantly a reflection of hydrothermal processes and iron in­put (Isley, 1995), with or without a biogenic influence (Cloud, 1973; Holm, 1987), then the repetitions may have nothing to do with annual fluctuations.

More recent interpretations of accumulation rates for banded-iron-formation rates in the Hamerslcy Basin and Transvaal deposits are based on detailed radiogenic chro­nology of sequences. They suggest depositional rates orders of magnitude slower than those proposed by Trendal and Blackley (1970), -0.001-0.004 mm a-I (Arndt et aI., 1991; Barton et al., 1994). These rates apply to both shale and banded-iron-formation deposits; underlying dolomites may have been deposited an order of magnitude faster than those. Deposition might not have been continuous, so that indi­vidual bands might have been deposited quickly, followed by a depositional hiatus. Such hiatuses are not interpreted

to reflect unconformities, and all extraterrestrial material deposited in an entire time package should be in the se­quence, perhaps concentrated at grain boundaries. The es­sential point is that a wide range of possibilities for the overall depositional rate for banded iron formations exists, and 1 mm a-I is perhaps at the very high end. It is not pos­sible to clearly establish depositional rates for the specific Isua and the Akilia banded iron formations that we analyzed (next section); we can only suggest and use a range of rea­sonable possibilities.

Table 1 shows the calculation Ir[sED] in ppb for background infa!1 from equation (I), assuming ¢m = 40 x 109 g a-I (Love amI Brownlee, 1993); [Ir]ET = 480 x 10-9g g-I (= 480 ppb; chondritic) (Anders and Grevesse, 1989); EBA = 5.1 x 1018 cm2;

PSED = PBIF 3,3 g cm-3, based on average mineralogy of BIF; and varied inputs of sedimentation rate from 0.100 mm a-I to 0.001 a-I, and of greater extraterrestrial background flux from 300 to 10,000x the present rate.

The expected [1' abundances range from -0.003 ppb for very rapid depositional rates and low ancient fluxes to -I I ppb for very slow depositional rates (roughly that of Cretaceous-TertialY boundary clay, for instance) and high ancient fluxes.

4.3.5. Search for enhanced (!.'(lralerreSlrial influx a/si­derophile elements. Mojzsis and co-workers studied aque­ous sediments from the early Archean of southwestern Greenland for analysis for trace elements, including II' (Mojzsis, 1997; Mojzsis et al., 1997: Ryder and Mojzsis, 1998). These oldest terrestrial sediments might be expected to preserve a signal of higher flux, according to their pre­cise age correlation with the lunar bombardment record and their depositional rate. While the methods and results will be detailed elsewhere (Mojzsis and Ryder, 2000), we pro­vide a summary here.

The samples selected by Mojzsis and co-workers were early Archean banded-iron-formation enclaves from Akilia Island (the oldest currently-known sediment); banded iron fonnations, quartzite, and "control" granitic Amitsoq gneiss from Innersuartut Island just south of Akilia (Fig. 3); and banded iron formations from the Isua supracrustal belt (Table 2). We also analyzed the Gunflint Chert, a sample of Proterozoic banded iron formation. The samples were prepared and analyzed using neutron activation techniques

TABLE I. Calculated 11' (ppb) in sedimcnts, from background flux.

Flux limes present rate

Sed rate, 111111 a-I 300 1000 2000 10000

1.000 0.0003 0.0011 0.0023 0.0114 0.500 0.0007 0.0023 0.0046 0.0228 0.100 0.0034 0.0114 0.0228 0.1140 0.050 00068 0.0228 0.0456 0.2280 0.010 0,0342 0.1140 0.2280 1.1400 0.005 0.0684 0.2280 0.4560 2.2800 0.001 0,3420 l.l400 2.2800 11.4000

Ryder el al.: Hea>:v Bombardment oIlhe Earlh 485

50km

51"W

early Archean (3700·3900 Mal IIlsua Supracrustal 8elt (;;Amitsoq gneisses

Fig. 3. Generalized geological map of southern West Greenland.

at the Johnson Space Center (JSC). The samples were pre­pared mainly as crushed, cleaned, interior, roughly whole­rock particles. Approximately 100-200 mg of particles of each sample were encapsulated in pure quartz tubes for ir­radiation and y-ray counting. All samples were counted three times (-0.5 week, I week, and 3 weeks after irradiation); some were counted yet again a few weeks later to improve the precision (detection limit) for II'. Data were reduced using standard procedures at the NASA Johnson Space Center laboratory (D. Mittlefehldt, personal communica­tion), The detection limit obtained was 0.4-0.8 ppb (20) for [1' for all but the most Fe-rich samples, for which the detec­tion limit was closer to 2 or 3 ppb (20) (Table 2),

The data showed that the samples contained little detri­tal material, consistent with their thin-section characteris­tics, with incompatible-trace-element abundances not unlike previous analyses of banded iron fomlations and related rocks (e,g., Dymek and Klein, 1988). None of the samples investigated, including the -2.I-Ga Gunflint Chert, had 11' above its detection limit for that sample (Table 2). Clearly none of the material we analyzed was a rapid fallout simi­lar to the Cretaceous-Tertiary boundary clay, for which we would expect several ppb [1'. However, in terms of a greater background flux at the time of even the oldest (Akilia) iron­stones, our data are open to several interpretations. If the depositional rate is truly very rapid (tenths ofmm a-lor so), then even under the highest expected ancient meteoritic flux our data would not detect the expected Lr «0.1 ppb, in some cases «0. I ppb). However, if the depositional rate was ac­tually more similar to that of shales or carbonates, or even somewhat faster, then our data indicate that the flux at the

486 Origin of the Earth and Moon

TABLE 2. Neutron activation analyses of rocks from southwest Greenland.

FeO· Na20 La lr Cr Co Ni (%) (%) (ppm) (ppb) (ppm) (ppm) (ppm)

Akilia Island banded iron formations> 3.85 Ga ANU-92-197/I-A 5.8 0.027 0.52 <.4 I.I 4.8 49 ANU-92-197/1-B 6.4 0.034 0.51 <.27 1.5 4.5 33 ANU-92-197/2-A 7.4 0.051 1.36 <.5 1.6 5.1 26 ANU-92-197/2-B 7.4 0.044 1.97 <.3 1.6 5.0 33 ANU-92-197/3-AI 7.8 0.030 0.57 <.5 1.4 5.5 40 ANU-92-197/3-A2 5.2 0.025 0.52 <.4 1.1 3.9 29 ANU-92-197/3-B 9.0 0.030 0.66 <.5 4.1 5.8 42 ANU-92-197-X 20.1 0.D15 0.78 1.6 8.7 114

Innersuartuut banded iron formations >3.77 Ga SM/l55746-A 19.1 0.017 0.68 <.9 3.9 8.3 74 SM/l55746-B 23.8 0.018 0.66 <.6 4.7 10.2 81 SM/l55746-X 18.1 0.008 0.92 0.4 2.3 5.0 19 SM/155746-C 28.0 0.034 2.04 <.8 6.7 9.1 41 SM/155746-D 36.1 0.030 2.52 <.9 7.5 10.0 27 SM/171770-A 4.5 0.030 0.34 <.4 1.7 2.7 21 SM/171770-B 13.0 0.050 1.49 <.6 4.6 6.6 23 SM/l71770-X 14.8 0.031 1.48 4.4 8.7 55 SMlI7177I-A 70.2 0.129 1.40 <1.5 87.5 10.1 56 SM/17177I-B 54.3 0.242 1.70 <.6 50.0 11.9 58 SM/l7177I-X 51.9 0.301 2.56 63.9 11.7 66

{sua banded iron formations 3.77-3.80 Ga /3446-AI 54.5 0.002 1.05 <1.8 7.1 17.5 99 /3446-A2 53.7 0.002 0.63 <1.5 7.2 16.0 73 /3446-B I 52.1 0.004 0.62 <.7 5.8 16.1 64 /3446-B2 51.7 0.002 0.67 <1.1 6.5 16.1 76 /3446-CI 52.9 0.002 0.58 <.7 6.9 15.9 86 /3446-C2 51.7 0.002 0.70 <1.2 6.2 14.5 57 /3451-A 52.5 0.003 0.33 <2.1 6.3 13.7 53 /3451-B 51.5 0.002 0.45 <1.9 7.0 15.8 86

{sua Mt. - Isua banded iron formations SM178/248471 5.2 0.010 0.13 0.5 0.6 0 SMiGR/93/44 54.0 0.000 0.19 4.1 4.1

Isukasia - Isua banded iron formations SM/GR/96/8 69.5 0.002 2.07 152.9 38.1 80 SM/GR/96/9 48.9 0.006 1.90 198.7 20.7 68 SM/GR/96/1 55.5 0.006 7.43 8.9 27.8 166

Innersuartut Amitsoq orthogneiss > 3.77 Go SM/l55742 10.7 2.746 8.70 0.0 4.1 0 SM/l71773-A 2.0 2.872 9.47 <.7 3.8 4.0 <17 SM/l71773-B 1.6 2.998 13.38 <.8 3.1 3.0 <18 SM/171773-X 2.8 2.504 12.89 3.6 6.2 0

GU'!flint Chert -2.1 Ga GF7-A 4.3 0.011 . 0.77 <.41 0.9 3.2 <10 GF7-B 3.4 0.012 0.56 <.21 0.5 2.6 <II GF7-C 4.2 0.010 1.01 <.30 1.7 2.1 <13

A = saw-cut free, small pieces; B = saw-cut free, larger pieces; C = saw-cut enriched; D = mafic-enriched separate; X =

remainder, fines. See text for analytical information. • Total Fe as FeO.

Ryder et al.: Hea,y Bombardment of the Earth 487

time of their deposition was not of the order of thousands of ti mes the present flux. Clearly, at the present time we cannot provide a more definitive answer; both a better un­derstanding of banded-iron-formation deposition rates and more precise methods of analysis for Ir are desirable.

More precise analyses have been made for Ir and Pt in some banded-iron-formation samples from Akilia Island (Arnold et al., 1998; Anbar et aI., 2000). The analyses used a NiS fire assay and isotope dilution ICP-MS. Detection lim­its were -0.003 ppb Ir and -0.030 ppb Pt for the samples analyzed, which had abundances below those detection lim­its. This is somewhat surprising as the crustal background value is about 0.020 ppb for Jr. With such precision, even for a flux of 2000x the present and a sedimentation rate as fast as 0.5 mm a-I, Ir should have been detected in these samples (Table I). One can postulate even faster sedimen­tation rates, or nonrepresentative sampling (a nugget effect), or postdepositional loss of siderophile elements to explain these data. However, literal reading of the data would sug­gest that at the time of deposition, the bombardment rate was less than 2000x the present rate, probably much less.

Seventeen samples of Isua rocks, which could be up to 100 m.y. younger than the Akilia samples, were analyzed by Koeberl et al. (1998a,b, 2000) for their chemical com­position, including siderophile-element abundances. These authors also used Ni-sulfide + Te co-precipitation fire as­say and ICP-MS. The samples included metamorphic equivalents of turbidites, greywacke/felsic gneiss, conglom­erate/felsic metasomatites, pelagic shale, gravity flow from the Bouma sequence, phyllite, and banded iron formations. Four of 17 samples analyzed yielded measurable amounts of Ir (ranging from 0.06 to 0.18 ppb) above the detection limit (0.03 ppb), as well as Ru and Rh. The contents of the other siderophile elements (Pt, Pd, Au) are highly varied; chondrite-normalized abundance patterns show variations by a factor of 3-4 (Fig. 4). The elevated contents were observed in a variety of different rocks: one banded-iron-fornlation

.. 0.1 Ql u c: "0'" c: 0.01 ::> .0 <t: "0

'" 0.001.~ iii E 0 Z

0.0001

0.00001 Cr Co Ni Ru Pd Ir PI AJJ

Fig. 4. Chondrite-normalized platinum-group element (POE) abundance patterns in BIF (and other rocks) from Isua showing some Ir enrichment in BIF samples but a nonchondritic abundance pattern (after Koeberi el ai., 2000).

rock, one graywacke, one gravity flow sample, and one pelagic shale. Such variation could be the result of a terrig­enous detrital component. The elevated Ir content in the banded-iron-formation and pelagic shale samples may in­dicate a remnant meteoritic phase, but it is more likely they result from mafic contamination given the nonchondritic ratios of the other elements. If the Ir were demonstrably extraterrestrial, it would indicate a flux of -I 04x present for a deposition rate of 0.05 mm a-I, or a flux of -I 03x the present with more typical sedimentation rates. However, these samples are at least 50 m.y. younger than the Imbrium event, and therefore most likely postdate the main episode of late heavy bombardment.

The chemical search for an enhanced amount of extra­terrestrial matter in these Greenland samples was not deemed successful. This could indicate that either the rocks investigated were deposited very rapidly, that they do not overlap in time with the late heavy bombardment, or that the late heavy bombardment flux to the Earth was less in­tense than commonly predicted.

4.4. Search for Shocked Minerals

In normal terrestrial impact crater studies, the presence of shocked minerals is taken as confirming evidence for the impact origin of a purported astrobleme. The first petro­graphic search for shock features in rocks from Isua was reported by Koeherl and Sharpton (1988). Their study con­centrated on the search for shocked quartz; however, none was found. This is understandable given the multiple upper amphibolite-grade metamorphism that these rocks under­went after their formation; such metamorphism would have repeatedly annealed the quartz. On the other hand, a vari­ety of shocked minerals has been preserved in 2-Ga rocks from the Vredefort impact structure in South Africa. More recently, Koeherl et al. (I 998a,b) reported on a new search for shocked minerals in Isua rocks, this time using a min­erai that is more resilient in the face of recrystallization than quartz.

One of the best suited minerals for this purpose is zir­con, which has been demonstrated to record a range of shock-induced features at the optical and electron micro­scope level (e.g., Bohor et aI., 1993). Furthermore, zircon is very resistant to erosion and other forms of alteration, including high-grade metamorphism. While planar deforma­tion features in quartz may have long been annealed away, those in zircon have a good chance to survive for several billion years, as is indicated by the preservation of shocked zircons in rocks from the -2-Ga Vredefort and Sudbury impact structures. However, the identification of suitable early Archean rocks for such a study is difficult. Whereas sedimentary rocks containing detrital shocked grains would be best for this purpose, there is some controversy as to whether actual terrigenous clastic sediments occur at Isua (e.g., Rosing et al.. 1996). Koeherl et al. (I 998a,b) there­fore focused on some of the samples that have not been positively proven to be plutonic, and that have mixed zir­

488 Origin of the Earth and Moon

con populations either because they represent muitiphase intrusives, had an extended metamorphic history, or are eroded from a mixed source.

Several samples studied by Koeberl et a1. (l998a,b) yielded no zircons. Zircons were successfully separated from felsic schists, whose origin may be sedimentary. Grain mounts of hundreds of zircon crystals were studied; it was found that many grains are strongly fractured, but most frac­tures are of irregular shape or even of curved appearance. None of the crystals studied by Koeberl et at. (I 998a,b) showed any evidence of optically visible shock deforma­tion.

4.5. Search for Impact Debris

Recent progress has been made in examining the rock record for old cosmic spherules that would be a particularly enriched carrier for extraterrestrial signatures in sediments. Deutsch et at. (1998) found 18 magnetic spheres in a 5-kg sample of IAO-Ga red-bed sandstone from Finland. They assumed that all the spheres were extraterrestrial, and lim­ited their search to the 60-125-flm size fraction. Taylor and Brownlee (1991) discovered numerous micrometeorites in a Jurassic (190 Ma) hardground, and Taylor et at. (1996) analyzed a magnetic fraction from 2 kg of Oligocene sedi­ments and found about 250 cosmic spherules preserved. Given multiple estimates ofhigher impact fluxes in the early Archean and the obvious preservation of cosmic spherules even in very old rocks, the search for extraterrestrial sig­nals in the oldest sediments is of renewed interest for esti­mating past fluxes.

The oldest known terrestrial impact structures are the Proterozoic Vredefort and Sudbury structures, 2023 ± 4 and 1850 ± 3 Ma respectively (cf. Reimold and Gibson, 1996), which represent the complete documented pre-I.85-Ga ter­restrial impact record. Other evidence for early Archean impact events is less demonstrative. Some enigmatic spher­ule hOlizons at and near the contact between the ca. 3.5-Ga Fig Tree and Onverwacht Groups in the Barberton Moun­tain Land, South Africa, have been reported as possible impact ejecta horizons. These argumcnts are based on tex­tural features, enrichments in the platinum-group elements, near-chondri tic platinum-group-element patterns, and Ni­rich spinels. In the case of the Barberton spherule layers, Koeberl and Reimold (1995) argued that the platinum-group enrichment is not a primary feature of the spherulitic hori­zons, but rather is the product of secondary mineralization. None of these spherule layer samples contained any evi­dence for impact-characteristic shock-metamorphic defor­mation, such as planar deformation features in silicate minerals. Thus the impact origin of these spherule layers is debatable. On the other hand, a Late Archean (ca. 2.5 Ga) spherule layer from the Griqualand West Basin, South Af­rica, shows clear evidence of a primary meteoritical com­ponent (Koeberl et at., 1999; Simonson et at., 2000). How­ever, no similar spherule layers have yet been reported from any of the earliest Archean rocks.

5. IMPLICATIONS AND OUTLOOK

A search for any evidence on the Earth for traces of the late heavy bombardment is currently centered on petro­graphical and geochemical studies of the world's oldest supracrustal rocks in Greenland. Petrographic studies of zircons extracted from these rocks have so far failed to show evidence for shock metamorphism. Nor have any deposits with ejecta-like characteristics, such as spherule beds, been reported so far from field investigations. At the time of this writing, results of the chemical search for a meteoritic com­ponent in these earliest sedimentary rocks remain uncertain. Of the samples analyzed by us so far, only four samples of different composition from localities in Isua yielded IT abun­dances that are somewhat above the present-day background levels for crustal rocks. However, in chondrite-normalized abundance diagrams of the platinum-group elements, even these samples show nonchondritic patterns and probably represent contamination from mafic phases. In the absence of any sign of shock metamorphism or ejecta deposits, and with ambiguous geochemical signals, no direct and un­equivocal evidence of a late heavy bombardment on Earth can as yet be confirmed.

The possible reasons for the failure to obtain such di­rect evidence are manifold. First, the number of samples studied so far has been small. This is certainly the case for the search for shocked minerals, but the geochemical stud­ies should have fared better. Second, the samples chosen might not have been ideal for such a search. However, given the limitations of the early rock record, the samples stud­ied were among the best available. A petrographic search for shocked zircons should be extended to the detrital zir­cons of known age (3.8-4.0 Ga), and a statistically signifi­cant number of samples from different locations should be scanned. However, Cintala and Grieve (1998) suggested that very-large-scale impacts may yield higher relative amounts of melt and thus may not preserve much shocked material, in which case the absence of shocked zircons may not mean much. A heavy bombardment should include abundant smaller craters, quite capable of producing shocked mate­rials, including zircon, if the target is zircon-bearing.

Third, it is possible that the formation time of the rocks studied so far do not overlap with the period of late heavy bombardment of the Moon. While the zircons from the granitoid rocks that cross-cut the Akilia samples have been interpreted to be -3.85 Ga in age (Nutman et al., 1996, 1997), it has been argued by some workers that these zir­cons are inherited from a preexisting rock, and that the host gneisses in the Archean of West Greenland are themselves only about 3.65 Ga in age (Kamber and Moorbath, 1998), but this argument has been disputed on several counts (Mojzsis and Harrison, 1999, 2000). The metasedimentary rocks from Isua are most likely less than 3.80 Ga in age. In this case, the late heavy bombardment would have ceased, and no direct evidence for an extraterrestrial component could be obtained. Fourth, it may be that these Akilia sam­ples are indeed younger than the late heavy bombardment,

but only slightly so, such that the uncertainties in the ages of the two overlap. It could be that Imbrium and Orientale, and the basins and craters stratigraphically between them, are all very close to 3.86 Ga, and the Akilia samples are -3.84 Ga in age, and immediately postdate bombardment.

If so, the lack of a heavy bombardment signature in the latter could result from a very rapid decline in the heavy bom­bardment in the 3.86-3.85 Ga timeframe. This is certainly quite possible, and would have ramifications for the bom­

bardment history of the inner solar system and for the ori­gin of the population of impactors that the bombardment represents. Recent studies in celestial mechanics have led

to the proposal of a mechanism that could plausibly supply a short-time spike in an otherwise steady or decreasing background flux of impactors (Zappala et al., 1998); how­

ever, such sources nced to be quantified. Lastly, it is pos­sible - but not very likely - that the Akilia rocks predate intiation of bombardment at 3.9 Ga and therefore missed all the excitement (Anbar et aI., 2000).

Further studies will be necessary to clarify the timing of the late heavy bombardment on the Moon and its effects on the Earth, if indeed they are preserved here at all. The precise ages of the actual rocks studied, and available to be studied, need to be resolved. Furthermore, the ambiguity of using simple elemental concentrations of siderophile ele­ments suggests that specific isotope systems, such as Os (Koeberl and Shirey, 1997) or Cr (Shukolyukov and Lugmair, 1998), would be useful in future searches for evidence of

impacts on the early Earth.

AcknowledgmentIJ'. This research was supported by the Fonds zur Forderung der wissenschaftlichen Forschung in Austria (CK), by the Lunar and Planetary Institute (GR), and by the U.S. Na­tional Science Foundation (SJM). SJM acknowledges additional support from the NASA-supported UCLA Astrobiology Center, and the Danish Scientific Research Council through the Isua Multi­disciplinary Research Project directed by P. W. U. Appel. We ap­preciate comments on the manuscript by R. A. F. Grieve, E. Pierazzo, K. Righter, and an anonymous reviewer. The Lunar and Planetary Institute operates under contract NASW-4066 with the National Aeronautics and Space Administration. This paper is LPI Contribution No. 1003.

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