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 abun­
dance) 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 sedi­
ments 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 sedimen­
tation rate of the studied sediments was too fast and therefore too diluting to record an obvi­
ous 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
Collisions between planetary bodies have been funda­
mental in the evolution of the solar system. Studies under­
taken over the last few decades have convinced most
workers that the planets formed by collision and hierarchi­
cal growth starting from small objects, i.e., from dust to
planetesimals to planets (e.g., Wetherill. 1994; Taylor,
I992a,b; see also chapters in section II of this volume), and
not from condensation downward. Late during the accre­
tion of the Earth (some time after -4.5 Ga), when it had
reached about 70% of its eventual mass, it was most prob­
ably impacted by a Mars-sized or larger body (see Cameron,
2000). The consequences of such an impact event for the
proto-Earth would have been severe and seminal, ranging
from almost complete melting and formation of a magma
ocean, thermal loss of preexisting atmosphere, changes in
spin rate and spin-axis orientation, to accretion of material
from the impactor directly, or through rapid fall-out from
orbital debris below the Roche limit. Much of the material
blasted off in the impact eventually reimpacted the Earth;
some of the ensuing Earth-orbiting debris would have rap­
idly coalesced to form most of the Moon and probably some
smaller moonlets. Some of this geocentrically orbiting
material would have continued to impact both bodies for
perhaps tens of millions of years after the lunar forming
impact. The essential accretion and core formation of the
Earth appears to have been completed about 50-100 m.y.
after the initial collapse of the solar nebula (Lee and Halli­
day, 1995, 1996; Halliday et al.. 1996; Podosek and Ozima,
2000), in a timescale apparently more protracted than that
for smaller planetesimals and Mars. As a result of later
geological activity, no record of any primary accretion bom­
bardment history remains on the surface of the Earth.
Thc period on Earth between the end of accretion and
the production of the oldest known crustal rocks is com­
monly referred to as the Hadean Eon (Cloud, 1976. 1988;
Harland et aI., 1989; Taylor, 2000), which is a chronostratic
division (Fig. I). Its terminal boundary is actually not de­
fined on the Earth; Harland et af. (1989) equate it with the
Orientale impact on the Moon. Others do not even use the
teml Hadean, either distinguishing the chronometric divi­
sions of Archean Eon and (older) Priscoan time (Harland
et al., 1989), provisionally at 4.0 Ga. Here we use the term
Hadean to represent the time period between the formation
of the Moon at -4.5 Ga and the beginning of the continu­
ous terrestrial rock record at 3.8 Ga.
In contrast with the youthful age for the crust of the
Earth, the surface of the Moon displays abundant evidence
of an intense bombardment at some time between its origi­
nal crustal formation and the outpourings of lava that form
the dark mare plains. Even prior to the Apollo missions,
these plains were calculated to be about 3.6 Ga in age based
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).
Ryder et al.: Heavy Bombardmenl ol the Earth
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 specifi­
cally 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 sub­
jected to a significantly greater bombardment than was the
Moon, as it has a larger diameter and a much larger gravi­
tational 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 bombard­
ment 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 hydro­
sphere, 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, sug­
gesting 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, de­
scribe petrographical and geochemical attempts to investi­
gate if any coeval record has been preserved on the Earth.
2.
2.1.
THE BOMBARDMENT HISTORY OF
THE LUNAR HIGHLANDS
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 in­
clude 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 aver­
age 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 correct­
ness of Hartmann's conclusion was demonstrated by the
return of Apollo 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 ori­
gin, 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 high­
land 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 young­
est 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
General
Whereas there is almost no evidence for terrestrial wit­
nesses 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 felds­
pathic crust was essentially complete by -4.44 ± 0.02 Ga,
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
Ga
3.6 - , - - - - - - - - - r ­
3.7
3.8
3.9
** = age of oldest Akilia sediment
Apollo 11 mare lavas
Late
Imbrian
Apollo 17 mare lavas
!III!!l~!III!!lt:::::".~orientale basin (3.80/3.84?)
Schr6dinger basin
~~~.-.iL~
-"4_ _.lmbrium basin (3.85)
"
Serenitatis basin (3.89)
Crisium basin (3.89)
Nectaris basin (3.90/3.92?)
=:
-+-----------------------'---1--..
.~-??--. . -
4.0
.........~
.
4.1
/
/'A-P-O-II-0-1-5-A-p-e-n-n-in-e--'V
/
??
; .
4.2
/
/
4.3
/.?~
4.4
4.5
Numerous basins
South Pole-Aitken basin
1# jCProcellarum basin?
/
Bench KREEP basalts
(3.84)
{/
~rustal '-m-a-g-m-a-t-is-m--------'
and volcanism
1
Oldest ferroan anorthosites/
crustal formation
Magma ocean
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 Cri­
sium, 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 compa­
rable 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 dat­
ing of the boundaries from youngest to oldest, from the sim­
plest 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
Ryder et al.: Heavy Bombardment of the Earth
defined, and includes the mare basalts at the Apollo IS land­
ing 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 land­
ing 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 land­
ing 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 Wil­
helms, 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 cra­
ters (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:i­
dinger 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:i­
dinger 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 of Imbrium). The best way to date an im­
pact 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 Ar­
Ar systems (e.g., Bogard et aI., 1991). Recently, Haskin et
al. (1998) have argued that all Th-rich impact melt brec­
cias (low-K Fra Mauro) collected 011 the Apollo missions
are products of the Imbrium impact event. Despite the dat­
ing 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 forma­
tion 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
479
certainly no older than 3.870 ± 0.010 Ga, and probably no
older than 3.836 ± 0.016 Ga. Third, similar arguments ap­
plied to the contents of the Fra Mauro Formation and the
Cayley Formation suggest a similar age constraint. For ex­
ample, 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 Forma­
tion, but represent later local events (such as the 14310­
group 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) hav­
ing 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 Ba­
sin, 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 prob­
ably 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 pre­
cisely 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 con­
siderable 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 felds­
pathic granulites, that may well date back to the very earli­
est postaccretionary bombardment at about 4.4 Ga. Further­
more, 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 Pb­
isotope growth curve. There is also a lack of the comple­
ment 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 in­
tense activity terminated at 3.85 Ga with the near-simulta­
neous 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 su­
perposed 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 deter­
mined, 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 col­
lectable and recognizable.
The chronology outlined above suggests a massive de­
cline 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 fac­
tor of -4 x 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 forma­
tion 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 Nec­
taris 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 (in­
deed, 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 1000­
1500x that of the Phanerozoic, and in the period 3.85-3.80
Ga was probably at least 250-400x that of the Phanero­
zoic. 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 poten­
tial information gleaned from rare detrital zircon grains that
are up to 4.27 Ga in age, -94% of Earth history. As men­
tioned 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 abun­
dant, 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 basal­
tic in composition (e.g., Taylor, 1989. 1992. 1993; Arndt and
Chauvel. 1991). The existence of any substantial early feld­
spathic crust on the Earth seems precluded by the higher
pressure at shallower depth on the Earth (in contrast to the
Ryder et a/.: Heavy Bombardmenr of/he Earth
Moon, which does have an ancient feldspathic crust). Cal­
cium 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 pla­
gioclase crystallization. In addition, plagioclase itself can­
not 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/86S r 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 compos­
ite granitoid source of continental provenance for these zir­
cons (Maas et al., 1992; Mojzsis, 1998). Thus there is evi­
dence 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 ex­
isted on Earth during much of the Hadean Eon. The lack of
initial Hf-isotopic heterogeneity and the absence of nega­
tive 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). Gra­
nitic 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, I 997).
3.2. Effects of Ancient Impacting:
From Basins to Dust
Individual impacts have considerable physical (morpho­
logical) and chemical effects on the target and on the at­
mosphere. 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
481
spherules. The projectile and some target rock are vapor­
ized, and a fraction of the projectile vapor can be incorpo­
rated into melt and ejecta. Minerals of both the autoch­
thonous target and the allochthonous ejecta could exhibit
shock effects (e.g., planar deformation features, high-pres­
sure 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 me­
teorites 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 di­
ameters, 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 sec­
tion 2), very little of the Earth's surface in the period of
-3.9-3.8 Ga should have escaped being the target of sig­
nificant impacts at one time or another, and therefore es­
caped 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 en­
vironment.
Calculations that scale impact-melt production with in­
creasing crater dimension (Melosh, 1989; Cintala and Grieve.
1994, 1998) show a breakdown of this geometric relation­
ship for very large impact structures. As the magnitude of
the impact increases, the melt volume relative to the tran­
sient 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. There­
fore, 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 hun­
dred kilometers in diameter, thermal metamorphism may be
more important than shock metamorphism. However, cra­
ters 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 ef­
fects 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 activ­
ity, such as volcanism, proto-ocean basin formation, atmo­
spheric 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 diam­
eters 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 diam­
eters 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 con­
cluded 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-in­
duced 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 prob­
ably requires the impact of an asteroid several hundred ki­
lometers 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 blan­
kets, shock features, and input of siderophile elements. Si­
multaneously, there should be a correspondingly greater
abundance of input of interplanetary particles and continu­
ous 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.
4.1.
SEARCH FOR EVIDENCE OF A LATE
HEAVY BOMBARDMENT ON
THE EARLY EARTH
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 abun­
dant gneisses. The Itsaq Gneiss complex of West Greenland
is a 3000-krn 2 terrane dominated by orthogneisses of grani­
toid compositions that intrude, in some locations. packages
of associated sediments and volcanic rocks. These supra­
crustal rocks are composed of massive amphibolites and
complex metasomatic carbonates (metamorphosed equiva­
lents 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 rela­
tively 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 relation­
ships on Akilia island and the nearby islets of the Godthabs­
fjord 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 forma­
tions to the polyphase, geochemically heterogeneous ortho­
gneisses that intrude them. The geochronological relation­
ships 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, Moor­
bath 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 as­
similated preexisting rocks older than the intruding grani­
toid orthogneisses on Akilia and the surrounding islands.
However, evidence for much Pb contamination from hypo­
thetical 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 re­
spect 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 con­
trast to precise and concordant zircon geochronology (Mojz­
sis 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 quali­
fied 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 extraterres­
trial materials, using different techniques and samples from
both Isua and Akilia rocks from Greenland. Second, evi­
(;
Ryder et al.: Heavy Bombardment of the Earth
dence of detrital shocked minerals that might have formed
as a result of an inccssant early bombardment may be pre­
served. 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.
483
(low estimate) and -10 4 (high estimate) greater than at
present. While this estimate is based on visible lunar cra­
ters, generally of the order of a few kilometers in diameter
and larger, it is inferably true of smaller craters and of in­
terplanetary particles and dust as welL In a geochemical
sense, it does not matter whether a projectile makes a cra­
ter or bums up in the atmosphere; it wil\ be added to the
4.3. Meteoritic Siderophile-Element Signatures
sediment as the dust settles. On the Moon the extralunar
material also has high abundances of the siderophile ele­
4.3.1. Siderophile elements on the early Earth. The Earth ments (for example, the Serenitatis impactor was almost
is a highly differentiated body, with a core, a mantle, and certainly an EH chondrite, James, 1995). All lunar impact­
evolved crust. During planet formation, the highly sidero­ melt rocks from the late heavy bombardment contain Ir in
phile elements (e.g., Ir, Pt, Au) partition strongly into me­ the 2-20-ppb range (Papike et al., 1998). On Earth, a sedi­
tallic cores. The formation of the Earth's core was completed mentary layer at -3.85 Ga might show evidence for an in­
early, well before the formation of the most ancient of pre­ flux of siderophile elements from an enhanced continuous
servcd terrestrial crustal rocks, and certainly by the time of background fallout, or from a specific event comparable
lunar formation. Thus, Earth's earliest mantle and crustal with the Cretaceous-Tertiary boundary layer, where such
rocks were effectively stripped of their highly siderophile' events had a higher probability than at the present.
element inventory early on. However, the present-day up­
4.3.2. Terrestrial sources ofiridium in marine sediments.
per mantle has abundances of highly siderophile elements Experiments have shown that -50% of Ir in sediments is
much higher than expected from presently known silicate­ scavenged from seawater by Fe-Mn-O-OH particles (Anbar
metal distribution coefficients and under the assumption of et aI., 1996) in oxic to suboxic environments. Anoxic envi­
core-mantle equilibrium (lr -3 ppb) (Chou, 1978; Chou et ronments, such as would be the case for much of the
al., 1983; Newsom, 1990). The siderophile-element abun­ Archean hydrosphere (Holland, 1984), are not a major sink
dances show chondritic relative proportions, which plot for Ir because of the redissolution of particulate hydroxides,
subparallel to thc Clline. The addition of -0. 75% chondritic except at rapid sedimentation and relatively shallow water
material after tennination of the core-upper mantle equilib­ depths. Iridium is well mixed in the oceans: The residence
rium under increasingly oxidizing upper mantle conditions time for it in the hydrosphere is 2000-20,000 yr. This im­
seems necessary to explain the abundances and chondritic plies that extraterrestriallr could persist in seawater and be
relative proportions of the siderophile elements in the mantle incorporated into sediments by particulate scavenging be­
(e.g., Chou et al., 1983; Newsom, 1990; Holzheid and tween impacts of a frequency of Jess than 2000 yr. It was
Palme, 1998). The emplacement timing of such a veneer is also found by Anbar et al. (1996) that Ir (and as) abun­
not constrained by direct evidence, but is often invoked to dances in present-day seawater are extremely low. Thus
have been as early as 4.40 Ga, or as late as 3.80 Ga.
weathering and hydrothermal alteration of ultramafic rocks,
Siderophile elements are strongly fractionated during such as peridotite, which could supply Ir (and other plati­
partial melting; for example, basalts are strongly depleted num-group elements) to seawater, is insignificant in deter­
in Ir «0.05 ppb) relative to mantle peridotites (-3 ppb). In mining the abundances of these elements in present-day
rare circumstances, siderophile elements can be concen­ seawater.
trated in specific crustal reservoirs, e.g., platinum-rich lay­
Studies of REE distributions in banded iron formations
ers in some basic intrusions; these have relative platinum­ demonstrate that hydrothermal activity had a strong influ­
group element abundances that are strongly fractionated ence on overall seawater chemistry in the Archean (Bau and
from chondri tic values. More evolved rocks, such as Moller, 1993). The average concentrations of lr in pelagic
pyroclastics, granites, and the sediments derived from them, clays with sedimentation rates of -0.001-0.003 mm a-I
contain negligible siderophile-element abundances from range from 0.07 to 2.0 ppb (Barker and Anders, 1968; Kyte
terrestrial sources. The source of significant abundances of and Wasson, 1986); in metalliferous sediments that scavenge
siderophile elements in evolved crustal rocks, such as the Ir, concentrations are even higher (Anbar et aI., 1996). Some
melt rock at East Clearwater Lake crater (Palme et al., of these higher abundances might result from organic mat­
1979), or in the Cretaceous-Tertiary boundary clay layer ter scavenging, and therefore do not reflect extraterrestrial
(Alvarez et al., 1980), can be reliably attributed to an ex­ input directly. Because these pelagic sediments are very slow
traterrestrial source. Thus siderophile elements in sedimen­ to accumulate, they contain measurable lr even at the
tary rocks, other than in the rarest of circumstances, can be present-day very low rates of meteoritic input.
4.3.3. Extraterrestrial sources or iridium to the hydro­
taken as an indication of an extraterrestrial flux at the time
of formation of the sediments, particularly if they are in sphere. Estimates of the influx of extraterrestrial matter
chondritic relative abundances.
reaching Earth's surface during the past 100 m.y. have been
Estimates of the flux of extraterrestrial material to the the subject of numerous studies aimed at quantifying the
Earth, based on lunar stratigraphic-chronologic studies dis­ current rate of dust accretion and the composition and
cussed above, suggest that during the peak of the late heavy source of the materiaL A number of methods have been used
bombardment this flux was, or ranged, between -3 x 102 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 ob­
servations. 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 10 9 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 as­
sumption 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 abun­
dances 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 «I 0 ~m. These grains are
quickly incorporated into sediments (Esser and Turekian,
1988). The finer dust grains are probably sensitive to sea­
water oxidation and hydrolysis after burial and have prob­
ably 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 pre­
serve a signal of higher incident fluxes from interplanetary
dust particles, micrometeorites, local impacts, airburst,
cometary showers, and ablation products of such phenom­
ena. 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 de­
sirable, because they have a higher proportion of extrater­
restrial 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 neg­
ligible contributions from the weathering detritus of igne­
ous 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 dis­
trict of southern West Greenland, and also younger ones
from West Australia and southern Africa, contain no signifi­
cant 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-iron­
formation 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+(Fe 3+h
0(OH)6 transforms to Fe2+(Fe3+)204 (magnetite), which has
low solubility in seawater. In general, detailed sedimento­
logical 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 envi­
ronments 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 for­
mations are considered by many workers as being equiva­
lent to varves associated with seasonal changes in upwelling,
productivity, and local 0 production and other factors (Hol­
land, 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 iron­
stones, these authors estimated a deposition rate of 0.65­
1.3 mm a-I, which is much faster than even typical detrital
sediments such as shale and siltstone. However, while band­
ing 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 domi­
nantly a reflection of hydrothermal processes and iron in­
put (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 chro­
nology 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 indi­
vidual bands might have been deposited quickly, followed
by a depositional hiatus. Such hiatuses are not interpreted
Ryder el al.: Hea>:v Bombardment oIlhe Earlh
to reflect unconformities, and all extraterrestrial material
deposited in an entire time package should be in the se­
quence, perhaps concentrated at grain boundaries. The es­
sential 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 pos­
sible 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 rea­
sonable possibilities.
Table 1 shows the calculation Ir[sED] in ppb for background
infa!1 from equation (I), assuming ¢m = 40 x 10 9 g a-I (Love
amI Brownlee, 1993); [Ir]ET = 480 x 10-9 g g-I (= 480 ppb;
chondritic) (Anders and Grevesse, 1989); EBA = 5.1 x 10 18 cm 2 ;
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/si­
derophile elements. Mojzsis and co-workers studied aque­
ous 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 pre­
cise 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 pro­
vide 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,
1.000
0.500
0.100
0.050
0.010
0.005
0.001
111111
a-I
300
0.0003
0.0007
0.0034
00068
0,0342
0.0684
0,3420
1000
0.0011
0.0023
0.0114
0.0228
0.1140
0.2280
l.l400
2000
10000
0.0023
0.0046
0.0228
0.0456
0.2280
0.4560
2.2800
0.0114
0.0228
0.1140
0.2280
1.1400
2.2800
11.4000
485
50km
51"W
early Archean
( ; ; (3700·3900 Mal
I I l s u a Supracrustal 8elt
Amitsoq gneisses
Fig. 3.
Generalized geological map of southern West Greenland.
at the Johnson Space Center (JSC). The samples were pre­
pared mainly as crushed, cleaned, interior, roughly whole­
rock particles. Approximately 100-200 mg of particles of
each sample were encapsulated in pure quartz tubes for ir­
radiation 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 communica­
tion), The detection limit obtained was 0.4-0.8 ppb (20) for
[1' for all but the most Fe-rich samples, for which the detec­
tion limit was closer to 2 or 3 ppb (20) (Table 2),
The data showed that the samples contained little detri­
tal material, consistent with their thin-section characteris­
tics, 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 simi­
lar 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) iron­
stones, 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 ac­
tually 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.
FeO·
(%)
Neutron activation analyses of rocks from southwest Greenland.
Na 20
(%)
La
(ppm)
lr
(ppb)
Cr
(ppm)
Co
(ppm)
Ni
(ppm)
Akilia Island banded iron formations> 3.85 Ga
ANU-92-197/I-A
5.8
0.027
ANU-92-197/1-B
6.4
0.034
ANU-92-197/2-A
7.4
0.051
ANU-92-197/2-B
7.4
0.044
ANU-92-197/3-AI
7.8
0.030
ANU-92-197/3-A2
5.2
0.025
ANU-92-197/3-B
9.0
0.030
ANU-92-197-X
20.1
0.D15
0.52
0.51
1.36
1.97
0.57
0.52
0.66
0.78
<.4
<.27
<.5
<.3
<.5
<.4
<.5
I.I
1.5
1.6
1.6
1.4
1.1
4.1
1.6
4.8
4.5
5.1
5.0
5.5
3.9
5.8
8.7
49
33
26
33
40
29
42
114
Innersuartuut banded iron formations >3.77 Ga
SM/l55746-A
19.1
0.017
SM/l55746-B
23.8
0.018
SM/l55746-X
18.1
0.008
SM/155746-C
28.0
0.034
SM/155746-D
36.1
0.030
SM/171770-A
4.5
0.030
SM/171770-B
13.0
0.050
14.8
0.031
SM/l71770-X
SMlI7177I-A
70.2
0.129
SM/17177I-B
54.3
0.242
SM/l7177I-X
51.9
0.301
0.68
0.66
0.92
2.04
2.52
0.34
1.49
1.48
1.40
1.70
2.56
<.9
<.6
0.4
<.8
<.9
<.4
<.6
3.9
4.7
2.3
6.7
7.5
1.7
4.6
4.4
87.5
50.0
63.9
8.3
10.2
5.0
9.1
10.0
2.7
6.6
8.7
10.1
11.9
11.7
74
81
19
41
27
21
23
55
56
58
66
{sua banded iron formations 3.77-3.80 Ga
54.5
/3446-AI
0.002
/3446-A2
53.7
0.002
52.1
0.004
/3446-B I
0.002
/3446-B2
51.7
52.9
0.002
/3446-CI
0.002
/3446-C2
51.7
/3451-A
0.003
52.5
/3451-B
0.002
51.5
1.05
0.63
0.62
0.67
0.58
0.70
0.33
0.45
7.1
7.2
5.8
6.5
6.9
6.2
6.3
7.0
17.5
16.0
16.1
16.1
15.9
14.5
13.7
15.8
99
73
64
76
86
57
53
86
{sua Mt. - Isua banded iron formations
SM178/248471
5.2
0.010
SMiGR/93/44
54.0
0.000
0.13
0.19
0.5
4.1
0.6
4.1
0
Isukasia - Isua banded iron formations
69.5
SM/GR/96/8
0.002
48.9
SM/GR/96/9
0.006
SM/GR/96/1
55.5
0.006
2.07
1.90
7.43
152.9
198.7
8.9
38.1
20.7
27.8
80
68
166
Innersuartut Amitsoq orthogneiss > 3.77 Go
10.7
2.746
SM/l55742
2.0
2.872
SM/l71773-A
SM/l71773-B
1.6
2.998
2.8
2.504
SM/171773-X
8.70
9.47
13.38
12.89
<.7
<.8
0.0
3.8
3.1
3.6
4.1
4.0
3.0
6.2
0
<17
<18
0
<.41
<.21
<.30
0.9
0.5
1.7
3.2
2.6
2.1
<10
<II
<13
GU'!flint Chert -2.1 Ga
GF7-A
GF7-B
GF7-C
4.3
3.4
4.2
0.011
0.012
0.010
. 0.77
0.56
1.01
<1.5
<.6
<1.8
<1.5
<.7
<1.1
<.7
<1.2
<2.1
<1.9
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
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 un­
derstanding 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 lim­
its were -0.003 ppb Ir and -0.030 ppb Pt for the samples
analyzed, which had abundances below those detection lim­
its. 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 sedimen­
tation rates, or nonrepresentative sampling (a nugget effect),
or postdepositional loss of siderophile elements to explain
these data. However, literal reading of the data would sug­
gest that at the time of deposition, the bombardment rate
was less than 2000 x 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 com­
position, including siderophile-element abundances. These
authors also used Ni-sulfide + Te co-precipitation fire as­
say and ICP-MS. The samples included metamorphic
equivalents of turbidites, greywacke/felsic gneiss, conglom­
erate/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:
'c:"
"0
::>
0.01
.0
<t:
"0
'"
.~
iii
0.001
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).
487
rock, one graywacke, one gravity flow sample, and one
pelagic shale. Such variation could be the result of a terrig­
enous detrital component. The elevated Ir content in the
banded-iron-formation and pelagic shale samples may in­
dicate 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 extra­
terrestrial 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 in­
tense 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 petro­
graphic search for shock features in rocks from Isua was
reported by Koeherl and Sharpton (1988). Their study con­
centrated on the search for shocked quartz; however, none
was found. This is understandable given the multiple upper
amphibolite-grade metamorphism that these rocks under­
went after their formation; such metamorphism would have
repeatedly annealed the quartz. On the other hand, a vari­
ety 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 min­
erai that is more resilient in the face of recrystallization than
quartz.
One of the best suited minerals for this purpose is zir­
con, which has been demonstrated to record a range of
shock-induced features at the optical and electron micro­
scope 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 deforma­
tion 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) there­
fore 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 frac­
tures 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 deforma­
tion.
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 lim­
ited 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 sedi­
ments and found about 250 cosmic spherules preserved.
Given multiple estimates of higher impact fluxes in the early
Archean and the obvious preservation of cosmic spherules
even in very old rocks, the search for extraterrestrial sig­
nals in the oldest sediments is of renewed interest for esti­
mating 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 ter­
restrial impact record. Other evidence for early Archean
impact events is less demonstrative. Some enigmatic spher­
ule hOlizons at and near the contact between the ca. 3.5-Ga
Fig Tree and Onverwacht Groups in the Barberton Moun­
tain Land, South Africa, have been reported as possible
impact ejecta horizons. These argumcnts are based on tex­
tural features, enrichments in the platinum-group elements,
near-chondri tic platinum-group-element patterns, and Ni­
rich 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 hori­
zons, but rather is the product of secondary mineralization.
None of these spherule layer samples contained any evi­
dence for impact-characteristic shock-metamorphic defor­
mation, 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 Af­
rica, shows clear evidence of a primary meteoritical com­
ponent (Koeberl et at., 1999; Simonson et at., 2000). How­
ever, 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 petro­
graphical 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 com­
ponent 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 abun­
dances 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 un­
equivocal evidence of a late heavy bombardment on Earth
can as yet be confirmed.
The possible reasons for the failure to obtain such di­
rect 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 stud­
ies 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 stud­
ied were among the best available. A petrographic search
for shocked zircons should be extended to the detrital zir­
cons of known age (3.8-4.0 Ga), and a statistically signifi­
cant 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 mate­
rials, 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 zir­
cons 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 sam­
ples are indeed younger than the late heavy bombardment,
R.l'der et al.: Heavy Bombardment of the Earth
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 bom­
bardment 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 ori­
gin 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 pos­
sible - 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 ele­
ments 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. Na­
tional Science Foundation (SJM). SJM acknowledges additional
support from the NASA-supported UCLA Astrobiology Center,
and the Danish Scientific Research Council through the Isua Multi­
disciplinary Research Project directed by P. W. U. Appel. We ap­
preciate 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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