Tect0r10phvsics

Elsevier Publishing Company, Amsterdam - Printed in The Netherlands

THE BEGINNING OF CONTINENTAL EVOLUTION

CIIORGE W. WETHERILL

Departrmrt o.f f’1a’larwtar.v and Space Science. Utliwrsitj’ of’ California. LOS Arlgeles. Calif: (C~‘.S..-l.)

(Received August 25, 1971)

(Resubmitted November 2. 1971)

ABSTRACT

Wetherill, G.W., 1972. The beginning of continental evolution. In: A.R. Ritscma (Editor). The LIpper

Ma?ftle. Tectorzophysics, 13(1-4): 31-45.

Lunar studies carried out in connection with the Apollo program permit considerable understanding

of the early history of this planetary body. The resulting thermal history involves primary fractionation

of the outer half of the lunar mass by accretional energy followed by subsequent formation of the mare

basalts by radioactive heating at the boundary between the fractionated and unfractionated regions.

Application of this same model to the earth implies the entire earth was initially melted and geochem-

ically fractionated. Review of the geological and geochronological evidence for the most ancient rocks

indicates that not until 3400 m.y. ago did the earth cool sufficiently to permit the formation of ertrn-

sive areas of stable crust, and that this is the reason more ancient rocks have not been found more abun-

dantly.

This same line of reasoning implies that there is a class of planetary bodies intermediate in mass be-

tween the earth and the moon. possibly including Mars, which arc less magmatically active than either

the earth or moon, whereas the tectonic activity of others. possibly Venus, resembles that of the early

earth.

INTRODUCTION

Since the work of Patterson (1956) it has been known that the age of the earth (4600

f 100 million years) was considerably greater than that of the oldest (- 2800 m.y.) dated

terrestrial rocks (e.g. see compilation by Ahrens (1955)). Although at the time it seemed

quite possible that many older rocks would be found, this appears to be increasingly

unlikely.

During the last 15 years the development and widespread application of modern geochr J-

nological methods have opened the Precambrian to detailed investigation: the distinction

can now be made between the relatively young (e.g. 1000 m.y. old) granitic basement com-

plexes and the truly ancient 2800 m.y. old rocks which had previously been lumped to-

gether as “Archean”. With a single exception these much more extensive geochronological

surveys have succeeded in pushing back the age of the oldest rocks only to about 3400 m.y.

and the number of well-documented occurrences of that age is small. It therefore ap-

pears likely that the scarcity of older rocks is a fact which requires explanation.

Some possible hypotheses to explain this phenomenon are:

( 1 ) The age of the oldest rocks represent a “steady-state age”, as discussed by Wasserburg

(1961), and is merely indicative of the mean rate at which normal geological processes

erase the record of older events.

(3) The earth was formed as a cold planet and was heated slowly by long-lived radioactjv-

ity. About 3500 m.y. ago, melting began. The sinking of iron to the core released suffi-

cient additional energy to initiate the major chemical fractionation of the earth into trust.

mantle, and core. The oldest rocks of the crust necessarily postdate this event.

(3) The capture of the moon took place about 3500 m.y. ago. The close approach of the

earth and moon at that time resulted in the release of sufficient tidal energy in the earth’s

interior to destroy the record of previous earth history.

(4) The earth was formed at a high temperature and did not cool sufficiently to preserve

extensive stable crust until 3500 m.y. ago.

During the last 2 years, measurements made on sampies of rock returned from the

moon, together with geophysical measurements performed on the moon, have permitted

us to understand for the first time the early history of an observed planetary body. While

this understanding is not complete. it is sufficiently advanced to permit new insights into the

early history of the earth, and other planets of the solar system,

In this paper, I shall review the evidence obtained from terrestrial rocks, discuss the early

history of the moon, as revealed by lunar studies, and interpret the terrestrial data in the

Light of this new knowledge. This discussion will also lead to some inferences regarding

what one may reasonably expect to find on other planets.

THE OLDEST RUfKS Or: THE EARTH’S CRUST

On every continent there are large contiguous areas (* IO” km*) underlain by igneous

and metamorphic rocks greater than about 2400 m.y. in age. Typically, the age of these

rocks is 2600-2800 m,y. Altogether, these areas cover about 5% of the earth’s land surface

(Hurley and Rand, 1969). Within many of these areas, there are smaller regions in which

rocks as old as 3900 m-y. have been reported.

Although the term has sometimes been used in different ways, and consequently could

lead to misunderstanding, it is convenient to refer to these > 2400 m.y. areas as eratons.

There is a general belief that these cratons have become stabilized and are relatively in-

vulnerable to rejuvenation by younger tectonic events. This is certainly not always the case:

the 2700 m.y. rocks in Wyoming and Montana are in a region of the North American con-

tinent which has been profoundly affected by recent tectonic processes, presumably result.

ing from the overriding of an East Pacific plate by the North American-West Atlantic plate.

Furthermore, the presence of cratons on the margins of present-day continents should ex-

pose them to the tectonic processes associated with such margins: development of trenches

and subduction zones and calhsion of continental ~ithospher~~ masses. tether or not this

present exposure has characterized their past history is related to the question of whether

or not Laurasia and Gondwanaland existed as super-continents until their recent breakup

during the Mesozoic. The fact that cratons are frequently surrounded by not too much

younger (- 2000 m.y,) mobile belts, rather than by rocks of arbitrary age, suggests that at

Last in some cases they have been protected in some way from rejuvenation, Armour the

exact nature of this protection is at present an open and important question.

THE BEGINNING OF CONTINENTAL EVOLUTION 33

In a few cases the existence and age of > 3000 m.y. rocks within the cratons is fairly well

established; in other cases such certainty is not possible. At least in some cases the reported

occurrence of > 2800 m.y. old rocks has been shown to be incorrect. On the other hand, e\i-

dence is abundant that these older rocks have been affected by the later events which oc-

curred about 7-700 m.y. ago, and it is plausible that sometimes these later events have com-

pletely obscured the evidence for the older age of the rock.

Probably the best documented case for the existence and nature of rocks greater than

3000 m.y. in age is in the Barberton Mountain Land on the boundary between the Trans-

vaal and Swaziland in southern Africa. The oldest rocks comprise a greenstone terrane con-

sisting of low grade metavolcanic and metasedimentary rocks, which have many lithologic

and chemical similarities to greenstone belts found within cratons on other continents (An-

haeusser et al., 1969). The stratified rocks are divided into a lower Swaziland System, over.

lain by the upper Moodies System. The Swaziland System is in turn divided into the prin-

cipally volcanic Onverwacht Series, overlain by the predominantly sedimentary Figtree Se-

ries. These stratified rocks are surrounded and intruded by great masses of granitic rocks,

some of which have been named and correlated (Anhaeusser et al., 1969).

The 7920 m.y. age found for the post-Swaziland Kubuta pegmatite (Nicolaysen, 1954,

Aldrich et al., 1958)* supported the geological inference that these were indeed very ancient

rocks. The first clear evidence that most of these rocks were more than 3000 m.y. old W;IS

given by Allsopp et al, (1962). He obtained a Rb-Sr whole rock isochron age of 3070 m.y.

for granites designated by Hunter (1957) as X4”, i.e., a younger granite series intrusive

into post-Swaziland gneisses and migmatites. Although only four samples were measured,

they were highly enriched in radiogenic strontium. and fell very accurately on the isochron.

More recently, Oosthuyzen (1970) has made an extensive geochronological study of

granitic rocks from the Transvaal which post-date the Swaziland and Moodies Systems. Hi;

most interesting results are those obtained by isotopic U-Pb measurements on zircon, sphene.

and apatite. His data for two of these granites, Salisbury Kop and Dalmein. are shown in f‘ig. 1.

6-

0

. Solebury Kop lZlr,rcon Oalmem zircon

Saltsbury Kop III urco”

Dota from Oosthuyzen (1970)

4 8 12 16 20 24

p207,“235

1 1

Fig.1. Concordia plot of U-Pb measurements on minerals separated from - 3250 m.y. old granites ot

southern Africa.

l These, as well as all other Rb-Sr ages discussed in this paper, are calculated using a decay constant o;

1.39~10-“/ye31.

34 (i.\\‘. \\I.1 III KII I

As usual, these minerals yield discordant U- Pb ages. However, when the analyses are plot-

ted on a concordia diagram (Wetherill, 1956) they are most plausibly interpreted as - 3750

m.y. minerals which have lost lead at some subsequent time in their history. This implies

that the granites and the greenstone sequence intruded by the granites are at least this old,

and possibly significantly older. Oosthuyzen has obtained similar IJ -Pb ages for other gran-

ites in this region. Although the RbbSr data on these granites yield a somewhat younger

age (1990 m.y. ), presumably because of the effect of later events, Allsopp (I 961 ) has ob-

tained a high quality Rb-Sr whole rock isochron of 3700 m.y. on granite near Johannes-

burg, about 300 km to the east, thereby providing further evidence for - 3200 m.y. gran-

ites in this area.

The exact age of the older greenstone sequence is less clear. Van Niekerk and Burger

(1969) have published isotopic data for a single zircon separate from the quartz porphyry

phase of the Onverwacht Series, the lower series of the Swaziland System. The U---Pb ages

are discordant; the zo7Pb/‘06Pb age of 3360 m.y. is interpreted as the minimum age for the

formation of these volcanic rocks. Some further evidence for the existence of 3400-3500

m.y. rocks in this area has been based on the isotopic composition of lead separated from

sulfide minerals in the Onverwacht lavas (Van Niekerk and Burger, 1969), whole rock lead

from these lavas (Oosthuyzen, 1970) and lead from galenas from gold mines in the Barber-

ton region (Ulrych et al., 1967). These latter ages are strongly model-dependent, and conse-

quently it should be concluded that although it is quite likely that the greenstone sequence

is 3400-3500 m.y. old, this age remains to be definitely established.

Although the evidence is less clear. it is rather probable that portions of other cratons

are equally old. Data suggesting ages > 3000 m.y. have been reported for the Aldan craton

of Siberia (Rudnik et al.. 1969), the Fennoscandian shield of northern Europe (Sobotovich

et al., 1963). western Australia (Leggo et al.. 1965), India (Crawford, 1969) and Antarctica

(Halpern, 1970). In some cases these ages are based on measurements on lead minerals, and

are thus model-dependent: in other cases the experimental and geological uncertainties in

the determinations preclude definite assignment of these ages to the > 3000 m.y. range.

In none of these cases does the evidence approach the completeness of the African measure-

ments discussed previously.

The evidence for rocks of this age in North America is somewhat puzzling.

In the Canadian Shield there are many occurrences of greenstone belts resembling those

found on other continents and there has been a considerable quantity of geochronological

work done in these areas. particularly on the granitic rocks intruding these metasedimentary

and metavolcanic rocks. Despite considerable effort and careful study, none of these rocks

have been found to be older than about 7800 m.y. (Cast et al., 1958; Aldrich and Wetherill.

1960; Moore et al.. 1960; Osborne, 1964; Van Schmus, 1965: Wanless et al., 1966; Wanless

et al., 1967; Green et al., 1968; Purdy and York, 1968; Stockwell, 1968; Wanless et al., 1968.)

On the other hand Goldich et al. (1970) have reported ages of 3550 m.y. for granitic

gneisses in the Minnesbta River Valley, not obviously in a greenstone terrane, although the

exposure of the rocks is not good and their limited exposure might prevent observing the

entire geological picture. The best evidence for this great age is obtained by isotopic measure-

THL HtGINNING OE CONTINENTAL EVOLIITION 35

ments on zircons, which on a concordia diagram lie along a straight line between 1850 m.y.

and 3550 m.y. This high value for the lower intercept is itself not usual, on the other hand

it offers no firm basis for rejecting the data. Indeed it is what one might expect as a conse-

quence of the known events in this region of the continent about 1850 m.y. ago.

Although Rb --Sr whole rock ages up to 3800 m.y. are reported. these data are not con-

clusive. as they are in disagreement with one another, and fail to define an isochron.

Recently Black et al. (1971) have reported apparently convincing ages of 3980 + 170 m y.

in West Greenland.

Further work is needed in these areas to confirm or disprove this evidence for what

could prove to be the oldest surviving terrestrial rocks.

TIILRRIAL HISTORY Ot: THE MOON

During the last 7 years, studies of the moon carried out under the Apollo program have

provided a large quantity of information about the early history of that body. which in

turn permits us to infer considerably more about the early history of the earth. In particu.

lar, these considerations lead to the conclusion that the reason for the scarcity of terres-

trial rocks prior to about 3400 m.y. ago was a consequence of the earth’s being formed at

a high temperature, and not until about 1200 m.y. had elapsed did it cool sufficiently to

permit extensive preservation of stable lithospheric masses.

Measurement of the concentration of Rb and Sr and the Sr isotopic composition in lu-

nar samples has shown that a major fractionation of Rb relative to Sr took place on the

moon 4600 m.y. ago, i.e., immediately subsequent to the formation of the moon and the

solar system. First shown by Papanastassiou and Wasserburg (1970, 1971 ), this is clearly

seen by plotting the results of analyses of lunar soil samples on a Sr evolution diagram

(Fig.?) and noting that the data from the Apollo sites fall very near the straight line re-

presenting the 4600 m.y. isochron. The 87Rb/“6Sr ratio of these samples averages about

0.15, markedly higher than the value of this ratio for the sources of lunar basalts, which

average about a factor of 10 lower, as indicated by the initial 87Sr/86Sr ratio of the basal-

tic rocks. Further evidence for this early lunar differentiation has been obtained from

similar studies of highly differentiated lunar breccia, such as 120 13 (Asylum, 1970) and

the Apollo 14 rocks (Huneke et al., 1971) as well as U -Pb measurements (Tatsumoto,

1970; Silver, 1970; Gopalan et al., 1970). Taken altogether, these data show that the moon

underwent a major fractionation, immediately after its formation, enriching its surface re.

gions by factors of lo-100 in certain elements. No permissible concentration of long-lived

radioactivities could heat and differentiate the moon so quickly. As will become clear fro.11

the following discussion. the overall radioactivity of the moon can hardly be as great as

that of the chondritic meteorites or the sun, and as shown in Fig.3 chondritic radioactivity

concentrations would have required about 1500 m.y. to heat the moon to its melting point.

Some suggested heat sources, such a those involving short-lived 26Al, or electrical induction

(Sonett and Colburn, 1968) require an intensely active pre-main sequence sun occurring at

just the right time during planetary formation. There is no independent evidence nor cum-

712 46x._ ,.

.70 frothy aggregates

Lunar ftne surface maternal

i

700- 06990 (Basaltic Achondrite lnttial Value)

698* I 2 3

i

Rb67/Sre6

Fig.2. Sr evolution diagram for ApoIIo 11 and 12 soil fractions indicating primordial enrichment in Rb relative to Sr. Data from Cliff et al. (I971 and unpubIished).

“DIFFUSION DEPTH”(KM)

Approxlmote Mellmg

”

45 40 35 30 25

TIME BEFORE PRESENT ilO9 YRS)

Fig.3. Temperature increases produced by typical rocks with varying concentrations of radioactive eie- ments early in the history of the solar system. The upper abscissa labeled “‘diffusion depth” represents the characteristic depth of thermal diffusion associated with the time scale on the lower abscissa, e.g. an initially hot region at a depth of 100 km will remain hot until about 4.3~10~ years ago, a 200 km layer will not cool until about 3.4*I09 years ago. Also, a layer 50 km thick having the average radioac- tivity of mare basalt will be abte to sustain a steady-state temperature of only about 7OOp earty in lunar history, whereas a similar layer of alkahe basalt will be able to sustain a temperature above the melting point at its base.

THE BEGINNING OFCONTINENTAL EVOLUTION 31

pelling reason to believe that such sources were effective. Indeed, the absence of excess 26 Mg

in meteorites (Schramm et al., 1970) which were formed during the same epoch in solar

system history argues against this possibility. Furthermore, both of these sources would re-

sult in a moon which was hottest at the center, in conflict with measurements of the electri-

cal conductivity of the lunar interior (Sonett et al.. 1971), unless ad hoc assumptions are in-

troduced involving initial inhomogeneities in chemical composition or electrical conductiv-

ity. The most likely heat source for this initial differentiation of the moon is the gravitatio-

nal energy of lunar accretion.

In order for a sufficiently high temperature to be reached during this accretion, it is nec-

essary that the accretion of the moon occur on a time scale of - 10” years. Prior to the on-

set of melting. the characteristic heat diffusion length in silicate material corresponding to

this time interval is about 100 meters. Consequently. diffusion of heat into the lunar in-

terior will be negligible during the accretionary phase, and the initial temperature distribu.

fion will be inhomogeneous: hottest at the surface and coolest at the center.

The details of the accretion process are not known. During the earlier phases of accretiorr,

it is probable that the accretion rate is proportional to the surface area; later the effective

area of the moon will be augmented by its gravitational field. Toward the end of the accre-

tionary period, the supply of matter to be swept up will decrease. slowing the rate of accre-

tion: at the same time the colliding matter may well have significantly greater kinetic ener-

gy than that provided by the lunar gravitational field, thereby increasing the heat produc-

tion per unit accreted mass. Under the simplifying assumption that the rate of accretion is

proportional to the surface area, the initial temperature at a radius r from the center of the

moon will be given hy:

GM(r) ” T4 +c(T-- To) = y - PR

where: T = temperature at radius r; T,, = initial temperature of the accreting matter: u =

Stefans constant: r = total time required for accretion of the tnoon; p = density: R = final

radius of the moon; c = specific heat; G = gravitational constant;M = mass of moon out to

a radius r,

Substitution of numerical values into this expression leads to surface temperatures of

- 1500" K for accretion times of 1000 years. Once the melting point is reached. partial

melting and gravitational settling will probably mix the surface regions, resulting in approxr-

mately uniform temperature beyond the radius at which melting first occurs.

1000 years is a very short time in which to accrete the entire moon. However, it may not

be impossibly short. The time required depends critically on the density of fine grained

material available for accretion. If the present mass of the earth were distributed in the

region between 0.95 and 1 .OS a.u.. the density of matter in this region would be about

I O-'"g/cm", and the time required for the accrefion of the earth and moon would be quite

long: IO7 ~- I OS years. Under these circumstances. gravitational heating would be negligi-

ble. However, this accretional model may not be correct. There is no reason why the t-cgion

of the solar system presently occupied by the asteroids could not have initially been as

densely concentrated with planetesimal bodies as the vicinity of the earth. Following the for-

38 (s IL il’1.f Iii llll I

mation ot Jupiter, residual matter I‘rom the formation of Jupiter. asteroidal lllaterial beyc,llri

4 a.u., and material in the 2: 1 Kirkwood Gap would have been subjected to very strong pet.

turbations by Jupiter, and wt~~ld have presented the asteroid belt and inner solar systent

with an intense bombardment of meteoroida material in orbits similar to those of the

present meteoroids, but with about 10’ times the present intensity.

At the present time the lifetime for destruction of meter-size planetesimals is about I o7

years (Gault and Wedekind, 1969), whereas 10 km bodies should survive for about IO’

years. Under the hypothesized high flux following the formation of Jupiter, these lifetimes

will be decreased by a factor of I 07- 1 year and 100 years respectively. This implies

two things. First of aIll reasonably large embryos (i.e. IO km) must have formed in the

inner solar system if any further accretion is to be expected at all. Secondly, an enormous

mass of finely crushed material will be produced in the asteroid belt, which if it can be

brought into the region of’ the earth and moon. will greatly facilitate the rapid growth by

accretion of any embryos sufficiently large ttt survive the b~)Itibardment. The exact mecha-

nisms of this inward llli~r~tion of debris are not clear: it may be swept along by the volatile

material falling inward CO form the sun, its own viscous dissipation may provide an effective

mechanism for energy loss. In any case. the possible movement ot’hundreds of earth-masses

ot’such material through the inner solar system completely alters the previously mentioned

idea of simply sweeping up the mass of the earth and the moon from the present supply of

matter in this region of the solar system. Obviously, a key question in this whole problem

is that of the time and the rate of formation of jupiter. Until these problems are worked

out. it seems premature to say on a priori grounds whether gravitational accretion of the

earth and moon on a IO” year time scale is possible or not. It is also possible that a flux of

high velocity bodies during the terminal stages of lunar accretion could enhance the final

rate of energy pi-(~duction sign~i.ic~~Iltly. Bodies moving with a relative velocity of 25 km/

set would supply more than 100 times as much kinetic energy per unit mass as those having

an initial velocity small compared with the lunar escape velocity.

Available evidence does not permit firm statements regarding the plausibility of intense

high velocity bombardment early in lunar history. Rb $r measurements on Apollo 14 rocks

(Huneke et ai., 197 1) suggest that the Imbrmm basin was excavated by collision as recent-

ly as 3.Y*109 years ago. The Orientale basin is unlikely to be aider, and measurenlents of

the crater frequency on the Orientale ejecta blanket show it experienced a bombardment

flux 10 times as great as the Apollo 1 1 site in Mare Tranquillitatis (Gault, 1970) dated at

3.6. I O9 years (Papanastassiou et al., 1970). This implies a flux during this period in lunar

history declining with a half-life of about 10’ years. Extrapolating this back to the time of*

formation of the moon, 4.6*10g years ago, a terminal high energy flux about 1000 times

as great as is predicted, in agreement with crater frequencies in the Southern Highlands.

If these bodies were of cometary nature, important quantities of volatile material would be

supplied by them. but this flux is insufficient to produce significant heating, being only

about 10,’ 10” the present extraterrestrial flux per unit area on the earth’s surface, which

would be c~~nlparable to the present surface heat t‘lux from the earth’s interior. and much

less than the flux of solar energy.

Tilt- BFGINNING OF CONTINENTAL EVOLL’TION 39

However, it is not clear that this extrapolation is valid. If the bodies which formed Mare

Imbrium and Mare Orientale were in orbits extending into the asteroid belt with dynamic

lifetimes of about IO* years, one must also take into account their lifetime with regard to

destruction by collision with smaller bodies in similar orbits. Current estimates of the streng,h

of such bodies (Gault and Wedekind, 1969) together with theories of asteroidal collision

(Wetherill, 1967; Dohnanyi, 1969) suggest that during the first Sew hundred million years

of solar system history, the spatial density of these bodies could be great enough to permit

collisional destruction to predominate, resulting in much shorter half-lives, e.g. 1 O6 I O7

years, and consequently much greater terminal fluxes. Because of these uncertainties, the

role of high energy impacts during the terminal stages of lunar accretion must remain open.

There are two additional other important constraints on the thermal history of the moon.

The first is that at the present time the central portion of the moon is below the melting

point ~~ this is indicated by the measurement of Sonett et al. (1971) of the internal electri-

cal conductivity of the moon, and by its non-equilibrium figure. The other is that there

was still sufficient internal heat 3.25.109 years ago to produce the basaltic rocks at the

Apollo I7 site in Oceanus Procellarum. Incorporation of these two additional constraints

into a thermal model leads to the following thermal history of the moon (see Fig.4). This

Flux of ImpactIng bodies

LIP - P mixes sillclc materlal mto

/

I aw

melted region below

-I km SIIICIC crust enriched In K,U

2OO legIon melted during accretion, first depleted

km In K,U, later melting replenishes K,U from

i

below

50 region melted by long-lived radloactlvity

4 15 region rich In metallic Iron

I

undlfferentlated lunar material -never melted

1 1 1 To Center of Moon

I?g.4. Proposed thermal model of the moon, The outer - 200 km were differentiated at the time of

lunar formation. Subsequent radioactive decay melted the lower portion of the depleted repion. which

was to some extent replenished in radioactive elements, producing the high I( group of Apollo 1 1 ba-

salts. It is likely that metallic iron and other heavy minerals sank to the bottom of the region of the moon which was melted. The interior of the moon remains unmelted. and differs from average solar sys-

tem material insofar as chemical fractionation took place in the solar nebula and during accretion.

is essentially the history presented by Papanastassiou and Wasserburg ( 197 I ). and also dis-

cussed by others (Wood, unpublished; Hays, I97 I ).

(1) Accreting material low in radioactivity, i.e., potassium concentration sub-chondritic

by a factor of about 3.

(2) Accretional heating leads to melting at a radius of about 200 km from the present

(terminal) surface.

(3) Radioactive elements are enriched in liquid which floats on surface. At end of accre-

tional period melting extends to a depth of - 200 km. Bottom of melted zone contains

some fractionally melted material, with residual solids depleted in radioactivity.

(4) Radiation at surface freezes silicic melt as soon as flux of impacting bodies drops low

enough to permit it to be stable.

(5 j Melted region freezes after - 2.1 O8 years. Bottom of fractionally melted region still

remains at melting point after - 1 O9 years.

(6) Undifferentiated material heats slowly by radioactive heating, keeping top of this re-

gion at the melting point. Partially melted material. enriched in K, U migrates into zone

above. Eventually significant fraction of fractionally melted region, some portions of which

have been replenished in radioactive elements, is melted and comes to surface as mare ba-

salts.

(7) Decay of 23s U and 4o K ends process - 3.109 years ago. Interior, which was never hot,

fails to reach the melting point.

IMPLICATIONSWITH REGARDTOTHEE:ARLY HISTORYOFTHEEARTH

The proposed thermal history of the moon, invoking in an essential way gravitational

heating accompanying rapid accretion, has obvious implications with regard to the thermal

history of the earth.

If it is assumed that the earth and moon accreted in the same region of the solar system,

then it is possible to estimate the initial temperature distribution of the earth. Under this

assumption the time scale for the earth and the moon will be of the same order of magni-

tude. If the rate of accretion is proportional to the surface area, then the radius of the accret-

ing body will increase linearly, and the accretion of the earth will take four times as long

as that of the moon. However, the radius which is significant in accretion is the gravitational

radius rg :

where R is the physical radius; U, is the escape velocity; U is the geocentric velocity of the

accreting material. Because the escape velocity also increases with R, the accretion of the

earth will proceed very rapidly during its terminal stages, and it is therefore a satisfactory

assumption that the earth and moon accreted “competitively” in essentially the same period

of time.

Therefore melting of the earth would also take place when its radius was about 1500 km,

and eq.1 would then imply that material accreted at its final radius of 6400 km would be

at a temperature of about 20,000” K. Actually, it would probably not be this hot, because

TIIE RECINNING OFCONTINENTALEVOLUTION 41

once the melting point was reached mixing would be much more rapid, and the entire body

would tend to have a more uniform temperature of about 6000” K. These extremely high

temperatures would suffice to melt the earth throughout, and would not leave a central un-

differentiated region as was the case for the moon. Chemical differentiation would proceed

very rapidly, producting a surface region about 50 km thick highly enriched in radioactive

elements. As may be seen from Fig.3, this surface layer will be sufficiently thick to main-

tain a steady-state temperature of > 1500” at its base, assuming an average radioactive ele-

ment content equivalent to that of alkalic basalts. Furthermore. only about 10’ years would

be required to achieve this steady state. The actual temperature will be significantly higher

because of heat flow from the melted interior. As the surficial regions of the interior cools

to the melting point, the temperature in the radioactive layer will decrease. and the depth

at which melted rock occurs will become progressively deeper. After - 5.10’ years the loss

of heat sources due to radioactive decay will also become a significant factor in lowering

temperatures near the surface.

At the present time, the depth of incipient melting in the earth is about 100 km, and

this presumably is the depth at which the rocks are sufficiently weak to permit large scale

motions of lithospheric plates. Again, assuming the radioactive heat production of an alka-

lit basalt to represent a lithospheric average, the rate of heat production would be three

times as great as 3.5*109 years ago, and it seems likely that the depth of incipient melting

would have been about 30 km at that time.

Geological studies of the greenstone belts (Anhaeusser et al., 1969: Martin. 1 Y69) have

provided additional reasons for believing that these represent the remnants of small, thin,

and unstable lithospheric plates. Such thin plates would be especially vulnerable to rifting

and being destroyed as a consequence of being carried down into the region of melting by

subduction as well as by bombardment by planetesimals similar to these which produced

the lunar maria. The formation and destruction of such thin plates probably characterized

the surface of the earth during the first - 1 O9 years of earth history, after which time the

diminished heat production resulting from the decay of 4o K and 235 U permitted the growtlr

of sufficiently thick sialic blocks to more often survive these processes and remain “afloat”

while adjacent regions, analogous to the present ocean basins, participated more actively in

these tectonic processes.

Further evidence that the primary geochemical fractionation of the earth took place

early in its history is provided by the work of Hart and Brooks (1969) on the evolution of

the isotopic composition of Sr in the earth’s mantle. These measurements yield a 87 Sr/‘“Sr

ratio of 0.7017 + 0.0002 for 2700 m.y. old archean metavolcanics and 0.7076 + 0.0002

for modern basalts from oceanic ridges. Assuming, as is reasonable from meteoritic and lu-

nar studies, that the initial s7Sr/s6Sr ratio of the earth was essentially equal to that of the

basaltic achondrite meteorites. i.e. 0.6990, this implies that the R7Sr/86Sr ratio of the

source of the basalts increased by 0.0027 during the first 1900 m.y. of earth and only l/3

as much during the remaining 2700 m.y. While these results do not lead to a unique model

for the evolution of the mantle Rb/Sr ratio, they are consistent only with models in which

mantle depletion in Rb relative to Sr took place prior to the formation of the most abundant

ancient rocks, i.e. those of about 2700 m.y. in age.

In summary, the evidence is accumulating for the earth being very hot during the first

10’ years of its history. This conclusion is not entirely tied to a particular model for the

accretion of the earth and moon. For example, the alternative “precipitation hypothesis”

of lunar origin (Ringwood, 1970) also requires an earth which was initially melted through-

out. Therefore, although these ancient greenstone belts do not actually represent “original

crust”, they are the oldest crustal material which has been able to survive the active tecto-

nism of early earth history. From this point of view they come closest to the “vestige of a

beginning” that we are likely to find. More extensive and complete geological, geochemical,

and geochronological studies of these regions should permit us to go beyond uniformitarian-

ism and gain understanding of an earth quite different in many ways than the one we see

today. It is an interesting question as to whether or not all of these original crustal blocks

should be of nearly the same age. The present. admittedly fragmentary, evidence suggests

they are not, but may range from 3900 m.y. to 2800 m.y. in age. Whether or not this re-

presents a problem for the ideas discussed here is worthy of considerable thought.

EARLYHISTORYOFOTHERPLANETS

This same line of reasoning leads to conclusions regarding the tectonic development of

other planets.

Consider first planets formed at approximately the same distance from the sun as the

earth and moon, e.g. Mars and Venus. Planetary bodies significantly smaller than the moon

(e.g. 1300 km in radius) will undergo insufficient gravitational heating to reach the melting

point, and will heat slowly as a consequence of their approximately uniform concentration

of long-lived radioactivity. For lunar concentrations of radioactive elements this source will

fail to provide sufficient heat to differentiate the body: however, after - 1000 m.y. the tem-

perature could rise sufficiently to permit redistribution of strontium isotopes and loss of

inert gases. The absence of meteorites exhibiting this thermal history argues either that bod-

ies that have undergone this thermal history do not exist in our solar system, or that the

somewhat uncertain processes which place meteorites into earth-intersecting orbits have

failed to provide them to us. Heating by electrical induction should be most effective for

these smaller planets (Sonett and Colburn, 1968) and may have resulted in their complete

differentiation 4.6.10” years ago, in spite of their lower gravitational energy of accretion.

In planets larger than the moon, gravitational melting will involve an ever larger fraction

of the planetary mass, resulting in complete melting at a radius of about 2.500 km. For

such planets, the heat carried to the central region by sinking iron and high temperature

silicates will be sufficient to melt and chemically fractionate the previously undifferentiat-

ed core. Planets of this size will have their radioactivity concentrated in a surface layer

too thin to maintain steady-state melting beneath the surficial radioactive layer, and will

also be greatly depleted in internal sources of heat. Although the deep interiors of such

planets will still remain at or slightly below the melting point, the absence of internal heat

sources will preclude significant tectonic or magmatic activity following their initial differ-

entiation. Mars may be such a planet. Larger planets, such as the earth and Venus will

THF HEGINNING 01: CONTINENTAL FVOLllTION 43

have a surficial layer of radioactivity sufficiently thick to maintain high temperatures at its

base, which together with the increase in melting point due to pressure and the additional

gravitational energy released by differentiation, will continue to he tectonically active

throughout their history. The high surface temperatures of Venus of about 500” C may be

expected to bring the depth of incipient melting closer to the surface of that planet. It is

quite likely that the tectonic activity of Venus is still characterized by low relief and “thin

plate tectonics”, as was the case for primitive earth.

For planets accreting at greater distances from the sun it may be expected that accretion

times will be longer. According to the collision formula of opik ( 195 1). either in its original

form or modified for collisions of bodies in eccentric orbits (Wetherill. 1967), the collision

probability varies essentially as the inverse cube of the distance from the sun, and directly

proportional to the mass of material in orbit at that distance.

These circumstances will lead to very low initial temperatures for bodies forming in the

asteroid belt. unless some alternative heating mechanism was present. In the case of Jupiter,

the greater distance from the sun will be offset by the greater mass in orbit at that distance.

and accretion times for that planet should be sitnilar to the earth. resulting in estremelq

high initial temperatures for that planet. If the Calilean satellites of Jupiter accreted “com-

petitively” with Jupiter, thermal and tectonic histories similar to either the I~OOI~ or Mars

may be anticipated.

The smaller mass and greater distance of the other major planets will result in longer

accretion times and lower initial temperatures for these planets and their satellites. Smaller

planets, such as possibly Pluto and other undiscovered bodies in the outer solar system could

well require 1 OS years for their formation, and gravitational heating during accretion will bc

of negligible importance. At these great distances, electrical heating by the solar wind will

also be of minor importance. and radioactive heating will probably play the major role in

the differentiation of these bodies.

hlany of the statements of this paper are obviously speculative. However, it is important

to recognize the relevance of lunar and planetary studies to our understanding of the earth.

and it may be expected that future discoveries will permit discussions of this kind to be

made on a much firmer basis.

This work has been supported by NASA Grant NGL 05-007-005.

RE1:ERENCI.S

Ahrens. L.H., 1955. Oldest Rocks Exposed. Geol Sot. Aru.. S~JC’C. Pap., 62: 155m-168.

Aldrich, L.T. and Wetherill, G.W., 1960. Rb-Sr and K-A ages of rocks in Ontario and northern Minne- sota. J. Geoph.vs. Hes.. 65: 337-340.

.4ldrich, L.T.. Wetherill. G.W., Davis, G.L. and Tilton, G.R., 195X. Radioactive ages of micas from gran-

itic rocks by Rb-Sr and K-A methods. Trans. Am. Geophys. Uniorl. 39: 1124&l 134.

Allsopp, H.L., 1961. Rb .Sr age measurements on total rock and separated-mineral fractions l‘ronl the

Old Granite of the Central Transvaal. J. (;EoP/?.\‘s. Rrs.. 66: 1499-1508.

Allsopp. H.L.. Robcrtq. H.R., Schreincr. G.D.L. and IHunter. D.R.. 1962. Rb-Sr age mcasurcrnenta on

various Swaziland granites. J. Gcwpl7_vs. Rex. 67: 5307-53 13.

Anhaeusser. C.R., Mason, R., Viljoen. ll.J. and Viljoen. R.P., 1969. A reappraisal of some aspects of

Precambrian shield geology. Hull. (;l,ol. Sec. :I )x1 80: 2 175 -- 2200.

Arnold, J.R., 1965. The origin of meteorites as smail bodies, 3. General considerations. ~lstroplr~x J..

141: 1548-1556.

Asylum, L., 1970. Mineralogic and isotopic investigations on lunar rock 12013. Earrh Planet. Sci. Lerr 9: 137-- 163.

Black. L.P., Gale. N.H., Moorbath, S., Pankhurst, R.J. and McGregor, V.R.. 1971. Isotopic dating of

very early Precambrian amphibolite facies gneisses from the Godthaab District, West Greenland.

Earth Planet. Sri. Lett. ( 12: 245-259.

Cliff, R.A.. Lee-Hu, (‘. and Wetherill. G.W., 1971. Rb-Sr and U. Th-Pb measurements on Apollo 12

material. Proc. Apollo I-7 Lunar Sri Conf:, Suppl. 2. Geochim. Cosmochim. Acta, in press. Crawford. A.R., 1969. Reconnaissance Rb-Sr dating of the Precambrian rocks of southern peninsular

India. .1. G’rol. Sot. I!rdia. 10: 1 17 - 166.

Dohnanyi. J.S.. 1969. Collisional model of asteroids and their debris. J. GeophJls. Res.. 74: 2531-2.554.

Cast. P.W.. Kulp, J.L. and Long. L.I:., 1958. Absolute age of early Precambrian rocks in the Bighorn

Basin of Wyoming and Montana. and Southeastern Manitoba. Trans. Am. Geoph. Uniorl, 39: 322- 334.

Gault. D.E.. 1970. Saturation and equilibrium conditions for impact craters in the Lunar Surface. Radio Sci., 5: 213-291.

Gault. D.E. and Wedekind, J.A.. 1969. The destruction of tektites bymicrometeoroidimpacts.J. Geo-

p!r)x Res.. 74: 6780-6794.

Goldich. S.S., Hedge, C.E. and Stern, T.W.. 1970. Age of the Morton and Montevideo gneisses and re-

lated rocks, Southwestern Minnesota. Bull. Geol. Sot. .4m., 81: 3671-3695.

Gopalan, I<., Kaushal. S.. Lee-Hu, C. and Wetherill. G.W., 1970. Rb-Sr and U, Th-Pb ages of lunar

materials. Proc. Apollo I I Lunar Sci. Conj:. Suppl. 1.. Geochim. Cosmochim. Acta, 2: 1195-l 205.

Green, DC., Baadsgaard, H. and Gumming. G.L., 1968. Geochronology of the Yellowknife area, North-

west Territories. Canada. Can. J. Earth Sci.. 5: 725-735.

Halpern, M., 1970. Rb-Sr date of possibly 3 billion years for a granitic rock from Antarctica. Science,

169: 977.-978.

Hart, S.R. and Brooks, C.. 1969. Rb-Sr mantle evolutionary models. Carnegie Inst. Wash. Ann. Rep. Dep. Terr. Msg.. 1968-1969: 426-429.

Hays, J.F., 1971. Radioactive heat sources in the lunar interior. Preprint. Huneke. J.C., Papanastassiou, D.A., Podosek, F.A., Tera, F., Turner, G. and Wasserburg, G.J., 1971. Age

determinations on samples from the Apollo 14 landing site. Proc. Seattle COSPAR Meering. Space Research. XII, in press.

Hunter, D.R.. 1957. The geology. petrology, and classification of the Swaziland granites and gneisses.

Trams. Geol. Sot. S. Afi-., 60: 85-l 25.

Hurley, P.M. and Rand, J.R.. 1969. Pre-drift continental nuclei. Science. 164: 1229-1242.

Leggo. P.S., Compston, W. and Trendall, A.F., 1965. Radiometric ages of some Precambrian rocks from

the Northwest Division of Western Australia. J. Geol. Sot. Aust., 12: 53-65.

Martin, H., 1969. Problems of age relations and structure in some metamorphic belts of southern Africa.

&ol. Assoc. Carl. Spec. Pap., 5 : 17- 26. Moore, J.M.. Hart, S.R., Barnett, C.C. and Hurley, P.M., 1960. Potassium-argon ages in northern Mani-

toba. BUN. Geol. Sot. Am., 71: 225-230. Nicolaysen, L.O., 1954. Age Determinations otf African Pre-Cambrian Minerals. Thesis, Princeton Univ.,

Princeton, N.J., 101 pp., not published.

Oosthuyzen, E.J., 1970. The Geochronology of a Suite of Rocks from thegranitic terrain surrounding the Barberton Mountain Land. Thesis, Univ. Witwatersrand, Johannesburg, 110 pp., not published.

iipik, E.J.. 1951. Collision probabilities with the planets and the distribution of interplanetary matter. Proc. R Irish A cad., 54A: 164- 199.

THL. BFGINNING OI’CONTINENTAI. EVOI.UTION 45

Osborne. I:.F., 1964. Gcockro7zolo,q~ ir? Canada. R. SIX. Carl. Spec,. Puhl. h’o. 8. Ilniv. l‘ornnto Press,

Toronto, Ont., 156 pp.

Papanastassiou. D.A. and Wasserburp, C.J., 1970. Rb-Sr ages from the Ocean of Stormc. Earth Plawr. Sci. Lett.. 8: 269-218.

Papanastassiou. D.A. and Wasserburp, G.J., 1971. Lunar chronology and evolution from Rb- Sr studiey

of Apollo 1 I and 12 samples. Earth Planer. Sci. Letf.. 11 : 37-62.

Papanastassiou. D.A.. Wasserburg. G.J. and Burnrtt. D.S., 1970. RbbSr a:cs of lunar rocks from the Sea

of Tranquillity. Earth Planet. SC;. I.ett., 8: 1-l 9.

Patterson. C.C., 1956. Age of meteorites and the earth. Geocl7it77. Cosn7ochi77~. Acta. 10: 230-237.

Purdy. J.W. and York, D.. 1968. Rb-Sr whole-rock and I*;-Ar mineral ages of rocks from the Superior

Province near Kirkland Lake, northeastern Ontario, Canada. Can J. Earth Sci.. 5: 699%705.

Rinswood. A.E., 1970. Origin of the moon: The precipitation hypothesis. Earrh Plarlcr. SLY. l.~~tt.. 8: 131-140.

Rudnik. V.A., Sobotovich. E.B. and Terent’yev. V.M.. 1969. Archcan age of the oldest rocks of the Al-

dan comples. Dokl. Akad. Nauk S.S.S.R.. 188: 897-900. Schramm, D.N.. Tcra, I,. and Wasscrburg. G.J.. 1970. The isotopic composition of ‘hMg and limits on

‘“Al in the early solar system. Earth Plarler. Sci. Lcrt.. 10: 44-59. Silver, L.T., 1970. Uranium-thorium-lead isotopes in some Tranquillity Base samples and their implication

for lunar history, Proc. .Ipollo 1 I I.fofar Sci. Cmf:, Slrppl. 1.. Clieochirx Costnoclrinf. Acta, 2: 15 33-

1574.

Sobotovich, E.V.. Graschenko, S.M.. Alesandruk, V.M. and Shats, M.M.. 1963. Age determination of an-

cient rocks by lead isochron and spectral strontium methods. 1-1,. Akad. Yauk S.S.S.R.. Ser. Gtd.. IO: 3-14.

Sonctt. C.P. and Colburn. D.S., 1968. The principle of solar wind induced planetary dynamos. P/7),x

Earth Plarzer. Inter.. 1: 326-346. Sonett, C.P.. Smith. B.l:., Colburn, D.S.. Dyal, P.. Parkin, C.W.. Schubert, G. and Schwartz. I<.. 1971.

Lunar electrical conductivity profile. Nature. 230: 359-362.

Stockwell, C.H., 1968. Geochronology of stratified rocks of the Canadian Shield. Carl. J. Earrh Sci.. 5: 693m 698.

Tatsumoto. M., 1970. Age of the moon: An isotopic study of U - Th ~ Pb systematics of Apollo 11 lu

nar - sample5 - 11. Proc. .4pol/o 1 I Lfrrlar Sc,i. Co77f Suppl, I, Gc~~rhi777. Cosr77ochi777. .4 cfa. 2: 1595 -~

1612.

Ulrych, T.J.. Burger, A. and Nicolaysen. L.O.. 1967. Least radiogcnic terrestrial leads. b’arfll P/a/fct. Scf. l&t.. 2: 179-184.

Van Nirkerk. C.B. and Burger. A.J., 1969. A note on the minimum age of the acid lava of the Onvcr-

wacht Series of the Swaziland System. Trans. Geol. Sot. S. Af?.. 72: 9-21.

Van Schmus, R.. 1965. The geochronology of the Blind River ~ Bruce Mines Area. Ontario, Canada. J.

Gcol.. 75: 755-780.

Wanless. R.K., Stevens. R.D., Lachance. G.R. and Rimsaite. J.Y.H., 1966. Ape determinations and gee-

logical studies - K-Ar isotope apes, Report 6. Geol. Sum. Can.. Pap., 65-17. 101 pp.

Wanless. R.K.. Stevens, R.D.. Lachance. G.R. and Edwards, C.hl., 1967. Ape determinations and geolok-

ical studies: K-Ar isotope ages, Report 7. Geol. Surv. Can.. Pap., 66-17, 120 pp..

Wanless. R.K., Stevens, R.D., Lachance, C.R. and Edwards. C.51.. 1968. Age determinations and get’-

logical studies, I<-Ar isotopic ages, Report 8. Geol. Sftrv. Carl.. Pap.. 67-2. Parr ..l. 141 pp.

Wassrrburp, G.J.. 1961. Crustal history and Pre-Cambrian Time Scale. A77r7. :V. Y’. :lcaJ. Sci.. 91: 583-

594.

Wethcrill, G.\V.. 1956. Discordant uranium-lead ages I. Trans. An7. Geopl7.w Ut7ior7. 37: 320 326.

Wctherill. G.W.. 1967. Collisions in the asteroid belt. .I. G~oplrys. Kcs.. 72: 2429~-2444. Wetherill, G.W.. 1969. Relationships between orbits and sources of chondritic meteorites. In P. klillman

(Editor). Mercorite Research. Reidel. Dordrecht, pp. 573~-589.

Wood, John A., unpublished. Thermal history and early magmatism in the moon.