MAGAZINEMINERALOGICALVOLUME43

NUMBER

MARCH I979

325

Mineralogy of the planets: a voyage inspace and time
HALLIMONDSOCIETYLECTURE
OF

FOR

GREAT

CENTENARY
BRITAINAND

OF

MINERALOGICALIRELAND,

I976

JOSEPH V. SMITH
Department of the Geophysical Sciences, University of Chicago, Chicago, Illinois 60637, USAs u MMARY. Pertinent general properties of the planets arelisted. The condensation of the solar nebula is set in thecontext of stellar evolution and meteorites with sections on

astronomical observations, chemical composition of thesolar nebula, physical properties of the solar nebula, andchemical and physical aspects of condensation, accretion,and planetary differentiation. A cool nebula is preferred toallow survival of pre-solar grains with isotopic anomalies.
Equilibrium progressive condensation of the solar nebula
is regarded as a useful theoretical boundary, but complex
processes involving crystal-liquid differentiation in, andcollisions between, planetesimals are used to interpret theproperties of meteorites and terrestrial planets. Chemicaldifferentiation in the nebula begins with condensation andaggregation of dust, which can yield oxidized and reduced
products depending whether C/O is less or greater thanunity. Simple models for direct accretion of condensed
materials into planets are reviewed but not adopted.
Physical interactions involving small bodies include eollisional accretion of dust-covered bodies, and differentiation of silicate and metal from mechanical, magnetic, andelectrostatic forces. Physical and chemical differentiation
involving large bodies involves head-on and glancing
collision of planetesimals, orbital deflection, and disintegration within the Roche limit, and collision withdebris rings and moons. Planetary accretion: dynamics,
time scale, and heat sources involves more rapid growth ofa larger body than a smaller one with ultimate development of one planet in each feeding zone, which flares out
and ultimately overlaps with adjacent zones. Mars issmall, and a planet did not develop in the asteroid belt,
because of perturbations from Jupiter. The giant planets
deflected material into the inner solar system. Melting of
early planetesimals is invoked to explain differentiated meteorites. Chemical differentiation inside planet-

Copyright the Mineralogical Society

esimals and planets describes the phase equilibria formetal, sulphide, and peridotite, either dry, wet, or containing CO2. A wet body could begin crystal-liquid differentiation near I25o K with sinking of Fe,S-rich liquid andrising of basaltic melt. The peridotitic residuum mightundergo a subsequent differentiation at higher temperature under volatile-free conditions. Mineralogical storageof H20, CO2, S, C1, F, and alkalies is discussed. Chemicaldifferentiation in planetary atmospheres briefly mentions
escape of light species.
For the Earth, the early history is constrained byArchaean rocks dating from -3.8 x lO9 yr whose properties indicate a non-reducing atmosphere, and a mantle
that yielded volcanic rocks mostly similar to recent ones.
The upper mantle (above 200 km depth) contains peridotitic
rocks attributable to crystal-liquid differentiation andmetamorphism. Volatile elements exist in mica and otherminerals, but are sparse. Abundances of sider0Phile andchalcophile elements are high enough to require lateaccretion of material rich in these elements, the presenceof a barrier between upper mantle and core, and someextraction by sinking sulphide. The mantle (deeper than
200 kin) and core are inaccessible to direct study but
interpretation of seismic data coupled with high-pressurelaboratory studies requires inversion to dense phases in
the mantle (especially perovskite?) and presence of lightelements in the core (mainly S?). The bulk composition ismodelled by cosmochemical analogy constrained bygeophysical and geochemical parameters. The early condensate may be augmented by 1.5 +0.5? over (Mg+Si),
and metal by 1.2++O.I?, while alkalies are probably
depleted six-fold. Radial heterogeneity from a reduced
interior to an oxidized exterior is suggested. For theMoon, sections cover observations, petrologic interpretations, and bulk chemical composition. For the origin of the

J. V. S M I T H

Earth and Moon, age constraints, chemical constraints, anddynamical and accretional constraints allow comparison ofsuggested origins, with the conclusion that the Moon
formed by either fission or disintegrative capture during
early growth of the Earth, followed by simultaneous
accretion coupled with disintegrative capture ofplanetesimals.
Mercury must be Fe-rich, but the silicate may not be
just early condensate. For Venus, reviews are given o f thesurfitce properties, atmosphere, speculations on bulk composition and speculations on surface mineralogy and atmospheric compositions. The CO2 is in the atmosphere; most
H may have been lost with concomitant oxidation ofrocks; K/U ratios suggest basaltic and granitic rocks, andthe high-surface temperature (c.740 K) implies granulitic
metamorphism. For Mars, reviews are given of surface
morphology, atmosphere and volatiles, mineralogy andpetrology, and geophysical and geochemical models. Prolonged emission of Fe-rich lavas is suggested. Volatiles
were removed by mineralogical processes from the atmosphere to give ice caps and sediments affected by aeolian
processes and oxidation from photochemically generated
H20 2. Fe-rich layer silicates, maghemite, and Mg-sulphate may dominate the sediments.
The composition o f comets and interplanetary dust may
be inferrable from micrometeorites whose complex
properties are suggestive of carbonaceous meteorites.
Asteroids should be supplying at least many of themeteorites to Earth. A general description and review ofremote-sensing studies culminate in a review of spatialdescriptions and implications. The main belt is dominatedby dark C-type asteroids assumed to be the primordial
inhabitants produced by primary condensation and local
accretion. The inner zone contains some brighter S-type
asteroids interpreted as having undergone separation of
metal from silicate as in stony-iron meteorites, as well assome E- and M-types perhaps matching enstatite-bearing
meteorites and irons. Perhaps these differentiated(?)
asteroids, as well as basaltic Vesta, are strays perturbed
outwards from the inner solar system.
A review o f meteorites covers carbonaceous meteorites,

ordinary chondrites, enstatite chondrites and achondrites,
reduced irons, forsterite-bearing meteorites and silicate
inclusions in irons, irons, pallasites, chassignites andnakhlites, ureilites and lodranite, eucrites and shergottites,
diogenites, howardites and mesosiderites, oxygen isotopes,
and original location of meteorites. Emphasis is placed onmineralogical properties demonstrating crystal-liquid
differentiation, brecciation, agglomeration, and evenaqueous alteration and vapour transfer, though someevidence remains of direct condensation of gas to solid.
The meteorites demonstrate the existence of many parent
bodies (probably over 60), and most macrometeorites are
ascribed to collision debris from the main-belt asteroids
that has been deflected past Mars. The range of oxygenisotopes is explained by mixing of supernova debris richin ~60 with gas-solid differentiates of the nebula.
The culmination of the review is a suggested synthesis,
which emphasizes a new model of heterogeneous accretion from planetesimals whose diverse compositions
range from reduced material near Mercury to oxidizedmaterial from the asteroid zone outwards. Growth of aterrestrial planet begins from near-by slow planetesimals,
and ends with distant fast planetesimals. Only the Earthand Venus are big enough to retain debris from latevolatile-rich planetesimals deflected to high speed by the
giant planets. Mercury, Moon, and Mars are volatilepoor. The Earth is zoned from a reduced interior,composed of crystal-liquid segregates of planetesimals
captured early, to an oxidized exterior containing somematerial captured from outside the orbit of Mars. A
mantle barrier hindered chemical equilibration. Hydrogen loss augmented oxidation. The mantle may bechemically zoned inwards from olivine-rich composition
to pyroxene-rich composition, and the core should be
reduced and contain substantial S and perhaps C and P.
Mercury and Venus should have accreted reduced
material, and only Venus should contain volatiles
obtained from late planetesimals from outside theMartian orbit. Mars should have accreted mainly
S-type asteroids, and have captured little volatile-richmaterial.

Index

General properties of the planets
Condensation of the solar nebula
Astronomical observations
Chemical composition
Physical properties
Chemical and physical aspects of condensation
Chemical differentiation
Physical interaction o f small bodies
Physical and chemical differentiation, large
bodiesPlanetary accretion: dynamics, time scale, heatsources
Chemical differentiation inside bodies
Chemical differentiation in planetary atmospheres
Earth:
Early history
Upper mantle
Lower mantle and core
Bulk composition

page
4
8
9
9
o
II
II
14
15
17
2o
25
26
28
30
3I

page
Moon:
Observations
Petrologic interpretations
Bulk chemical composition
Origin of earth and moon:Age constraints
Chemical constraints
Dynamical and accretional constraints
Comparison of suggested origins
Mercury
Venus:
Surface properties
Atmosphere
Speculations on bulk composition
Speculations on surface mineralogy and atmospheric composition
Mars:
Surface morphology
Atmosphere and volatiles
Mineralogy and petrology

36
37
39
4o
4~
4o
4l
42
44
45
45
48
50
5:

M I N E R A L O G Y OF THE PLANETS
Geophysical and geochemical models
Phobos and Deimos
Comets, asteroids, and meteorites
Comets and interplanetary dust
Asteroids:
General description
Remote-sensing studies
Spatial descriptions and implications
Meteorites
Carbonaceous chondrites
Ordinary chondrites
Enstatite chondrites and achondrites: reduced
irons
Forsterite-bearingmeteorites and silicate inclusions in irons

52
54
55
55
56
57
59
6o
6o
62
64
66

THE Mineralogical Society of Great Britain and
Ireland has a distinguished record in promoting
research in geochemistry, mineralogy, and petrology over the past century, and I was delighted but
awed to accept the invitation to talk at thisCentenary on some topic that looks forward to thenext century. Predicting the future is uncertain
whether it relates to science (recall the complacencyof classical physicists at the end of the nineteenth
century), to personal matters (as Ecclesiastes
reminds us), or to public affairs.
Direct exploration of the mineralogy of the solar
system began with the manned and unmanned
missions of the last two decades, four centuries
after a natural philosopher in Padua, Galileo
Galilei, used his 3z-power telescope to discover theroughness of the Moon, the Medicean satellites ofJupiter, the phases of Venus, and the spots on theSun. Fearful of revolutionary ideas the Roman
church placed Galileo under house arrest, but John
Milton visited him and was undoubtedly stimulated in writing Paradise Lost: e.g. lines 563-75 in
Book VII.
The Planets in thir station list'ning stood,
While the bright Pomp ascended jubilant.
'Open, ye everlasting Gates', they sung,'Open, ye Heav'ns, your living doors, let in
The great Creator from his work return'dMagnificent, his Six days' work, a World;
Open, and henceforth oft; for God will deign
To visit oft the dwellings of just Men
Delighted, and with frequent intercourse
Thither will send his winged Messengers
On errands of supernal Grace.'
In spite of numerous historical records ofmeteorite falls, a century of scientific scepticism
about these messengers from the gods elapsed
before the nineteenth-century exploration of themineralogical mysteries of meteorites, typified by
the microscopic studies of Henry Sorby in Edin-

Irons
Pallasites
Chassignites and nakhlites
AngriteUreilites and lodranite
Eucrites and shergottites
Diogenites
Howardites and mesosiderites
Oxygen isotopes in meteorites
Original location of meteorites
Suggested synthesis:
Pre-planetary stage
Earth
Other terrestrial planets
Conclusion

3
67
7~
7I
7I
72
73
74
74
75
76
77
78
8o
81

burgh and Gustav Tschermak in Vienna (Wood,t968). [Milton's winged messengers would becomets not meteorites.]
Voyages to the Moon were envisaged at least two
millenia ago (Braun and Ordway, t975), eventhough cultural memory of the Sun and planets as
inviolable gods persisted strongly into Elizabethan
times: indeed this mythical inheritance still remains
in astrology, and may even have produced psychological resistance of scientists to melting of the
poetical pale and cold Moon. A Parisian lawyer,
Jules Verne, extrapolated current technology,
including that of ballooning, in his fanciful but
scientifically based adventure tales La Voyage au
Centre de la Terre in 1864 and De la Terra fi la Lunein I865. Here in London, an unsuccessful teacher,
H. G. Wells, made his fortune with scientificfantasies including The War of the Worlds in I898
and The First Men in the Moon in I9oI. Sciencefiction writers continue to speculate on explorationof solar systems, but leadership is passing toengineers and physicists whose plans include establishment of human communities on artificial satellites (e.g. O'Neill, I974, I975) and planetary surfaces. A session on 'Utilization o f lunar materials
and expertise for large scale operations in space' on
I6 March I976, at NASA-Johnson Space Center,demonstrated the difficulties of construction ofhuman dwellings in space. It is easy to scoff atsuggestions for capturing comets and asteroids,
and for wholesale modifications of the atmosphereof Venus to lower the temperature enough forhuman habitation but, when account is taken of theachievements of technology in relation to thencurrent knowledge (e.g. the building of Neolithic
temple-observatories, the Great Wall of China,
nuclear reactors, and space rockets) and of thepossibility of new scientific discoveries, it is perhaps
safer to believe that some suggestions will ultimately bear fruit.
Although space exploration has been one of the

4

J . V . SMITH

major forces underlying the revolution in the earthsciences, and although practical needs will require
even more intense study of the Earth from bothobservational and theoretical viewpoints, it isperhaps optimistic to hope that scientific exploration of the solar system will receive high priority
over the next century as unconfined populations
exacerbate the mad passions of religious, political,
social, and economic tensions. But what a magnificent dream would be the conversion of money
spent on dreadful war machines and dangerous
drugs into exploration of the near-by planets. Even
if the Earth were spared human catastrophes,establishment of manned laboratories on planets orasteroids is unlikely in the next half-century (andperhaps century), and scientific study probably will
be based largely on remote observations, except for
meteorites and lunar samples. Small samples from
Mars may be returned to Earth within 2o years.
Establishment of a scientific laboratory on theMoon may result as an ancillary to engineeringdevelopments.
Study of lunar samples and meteorites has
shown how difficult is the detective work even forsamples dissected and probed by the most sophisticated instruments. Shock and thermal metamorphism have removed evidence in many samples, andwe must accept that certain questions cannot beanswered with certainty. Even adherence toChamberlin's method of multiple working hypotheses may not lead to a definitive answer, andhumility rather than cocksure assertion is preferable in facing such complex questions as theemplacement of the Moon around the Earth.
Looking backwards, the Earth begins to lose its
pre-eminence because of the absence of rocks older
than - 3 . 8 x io 9 yr. Early hopes that the Moon
would be a cold body with undisturbed relics of theearly solar system were thwarted by the discovery
of extensive melting and brecciation. Mineralogicaland chemical studies of meteorites, and remotesensing of asteroids, are demonstrating the greatcomplexity of products resulting from chemicaldifferentiation of the solar nebula. Probably theasteroids and comets preserve samples of most if
not all stages of development of the solar nebula.
Astronomical observations of stars and interstellar
clouds should provide boundary conditions fortheories of the solar system. As orbiting telescopes
are developed with facilities for observation without loss of spectral range in the Earth's atmosphere,
many ideas of the mineralogy of the surfaces ofsolar-system bodies should be tested.
Therefrom, the mineralogy of the planets must
be envisaged in a very wide framework embracing
many aspects of astronomy, geochemistry, andgeophysics. Perhaps no scientist can hope to understand all the relevant factors, and all of us must
recognize the limitations of our competence. With
great hesitation I try to assemble ideas from many
disciplines but recognize that the present view ofthe mineralogy of the planets is highly imperfect.
Let future scientists take joy in their success inreaching a higher level of understanding than thatattained here.
The textbooks by Hartmann (1972), Kaula(i968), Haymes (I97I), Short (~975), and King(i976) cover many pertinent properties of theplanets, but are becoming obsolescent as new
discoveries appear. The September I975 issue of
Scientific American contains useful surveys.
For brevity, referencing is less detailed forgeneral concepts than for mineralogical ones. The
lecture has been revised to include publications up
to June I978.

GENERAL PROPERTIES OF THE PLANETSThe heavens themselves, the planets, and this centre,
Observe degree, priority, and place.
WILLIAMSHAKESPEARE,Troilus and Cressida, I. iii. 85
Table I lists pertinent physical properties of theplanets, which bear no simple relation to the
distance from the sun. Properties of the outerplanets and especially of their satellites are difficult
to measure from telescopes (Newburn and Gulkis,
I973; Morrison and Cruikshank, I974; Scientific
American, Sept. I975, PP. 26 and I52).
For massive planets the real density must be
adjusted for the compression before being used tolimit chemical composition. The estimated density
at lO kb and c.3oo K depends on the variation ofpressure, temperature, chemistry, and crystal struc-ture with depth, but all reasonable distributions
should give a density within c.o. I g/cm 3 of the listed
value. [Note that I bar = IO5 N/m 2 = Io 5 Pa.]
The pressure at the centre of a planet can be
estimated accurately only when the density distribution is known. The gravitational pressure P isrelated to the density p(r) by dP/dr = -GM(r)p(r)/r z, where M(r) is the mass from the centre
to radius r, and G is the gravitational constant. The density distribution can be determined from inversion of seismic data, as typified bydensity models obtained from free oscillations ofMINERALOGY

OF THE PLANETS

5TABLE I. Selected physical properties of planets
PlanetMass
I027 gRadius
1 0 3 kmReal
densityg/cm 3

Estimated
densityat IGPag/cm 3

EstimatedcentralpressurelOl~ Pa

Escapevelocity
km/secT ofsolid
surface
K

Distance
to SunlO8 kmMercury
Venus
Earth(Moon) .
Mars
Jupiter
SaturnUranus
Neptune
Pluto

0.33
4.87
5.98
o.o7
0.65
19oo
57o
87

2.44
6.05
6.38
1.74
3.39
7o
6o
26
25
3??

5.44
5.27
5.52
3.34
3.96
1.3
0.7

5-3
3.9
4.o
3.4
3.9

< 4-4
< 3~
37
o.47
< 3.6

4
IO
Ii
2
5
c.6o
c.4o

IOO-65o74o_+ IO
23o-32oxoo 40o15o 3oo
---

0.58
1.o8
1.5o
1.5o
2.28
7.78
I4. 3

1,2

IOO

o.7?

--

c.IO 4---c.6 x lO3
c.4 X I 0 3

I. 7

--

c.6 X IO 3

3??

--

?

C.20

--

28. 7

C.25

--

45.0

?

--

59

Data from various sources.
0

3485

fl

i

5000
I

1',

6

I

-

~_,

i

-

r(km)

.4

.2

o

P(lO"Po)

.1

't

'l',

I =

Gilbert & Dziewonski(1975) model I066B

0 =Feo.95Nio.05-E

9 b = Feo. 9 S~

74.100 ~-..~.N6
,

,

0.2

-

2

-

o

637l
I'

6350

~ 1 5 . 5 6Ahrens(1977)
3

"~ 5
3

'

Moss Fractions
inner core 0.0162
outer core 0.3089
mantle0.6699crust
0.0050"-.-V!...........

7 --

I

5700

13~,-~ o/ 1216203077
i~',.j12.057"I1~--\ "*" 10.92
I \t- "r9.977
I
9-I
b

-~

5951

i

1217
4000
3 2 1.5 1 .8

[

,

0.4

1

(r/r ~ )3

3.293

3.~?~

o.8

__

,.o

FIG. I. Density-depth profile of Earth in model io66B ofGilbert and Dziewonski (i975). Use of third power ofradius r (normalized to r o = 6371 km) results in linear
segments. See Ahrens (i978) for discussion of accuracy ofshock-wave data in curves a and b, respectively from
McQueen and Marsh (I966) and Ahrens (I976).
the Earth (e.g. fig. I based on Gilbert and Dziewonski, 1975). F o r the Earth a uniform density would
give a central pressure of 17 x IO t~ Pa (I.7 x I O 6bar) compared to c.37 x lO l~ Pa for models with an
Fe-rich core (fig. I). The density profiles ofMercury, Venus, and Mars are unknown, but thecentral pressures are unlikely to be greater and may
be less than the values listed in Table I, whichpertain to models with a heavy Fe-rich core.
Accurate determination of the moment of inertia
would provide a strong control on the degree ofchemical differentiation of these planets.

The higher is the escape velocity E = (2GM/R)89
the more difficult is loss of material by volatilization (M and R are the planetary mass and radius),
and the easier is retention of impact debris during
accretion. Modelling of impacts responsible for thelarge basins on the Moon, Mercury, and Mars isdifficult (e.g. Baldwin, 1963; French and Short,
I968; M o o r e et al., 1975; O'Keefe and Ahrens,
I977a, b) but probably most impact debris isretained on planets with R > 2ooo km for impact
velocities < IO km/s.
The temperature at the solid surface depends on
the heat flow from inside, the heat flow from theSun, heat transfer by an atmosphere, and the heatloss to space. Atmosphere-free bodies must experience a strong range of surface temperature in
response to variation o f heat inflow from the Sun
during rotation. Because of thermal resistivity ofsurface rocks the temperature wave at the surface isdamped within a few centimetres for rotation
periods up to 1 0 2 days, which covers all known
stony bodies. F o r an atmosphere-free body rotating with its equator in the plane of the ecliptic, thesmoothed surface temperature T(W) varies withlatitude 9 as A + T(o)cos"qJ where n ranges from 88
to 89depending on surface roughness (Bastin andButt, I975) and A, the equilibrium temperature formaterial unheated by the Sun, depends on theinternal heat flow and the galactic temperature. O n
the M o o n the temperature variations should be
mainly between 4oo and IOO K and on Mercury
between 60o and IOO K, though somewhat lower
values might occur in shaded spots near the poles.
The heavy atmosphere and fairly smooth surface ofVenus should lead to little temperature variation at
the surface ( < 3o K ?). Mars has complex variationsof surface temperature with latitude, elevation, andseason.
The variation of temperature with depth below

J. V. SMITH
the surface is poorly known for the terrestrial
planets, and is model-dependent. Perturbation of
the solar wind by the Moon leads to estimates ofthe electrical conductivity with depth, and theelectrical conductivity can be transformed to temperature, assuming that the Moon is composed ofeither Mg-rich olivine or orthopyroxene (Dyal,
Parkin, and Daily, i975; Sonett and Duba, I975).
Particularly important is the assumption of complete divalency of the iron, which results in a lower
electrical conductivity for a given temperature thanfor the assumption of some trivalency. Seismic data
indicate lower Q near the Moon's centre than forthe outer part, but presence of melt is disputable, as
indeed would be the chemical composition of any
inferred melt. Because of the evidence for early
crystal-liquid differentiation of the Moon under
dry conditions, the pressure-temperature curve(denoted prete) for at least the outer part of theearly Moon probably became near to that for the
beginning of melting of dry peridotite. The exact
position of the curve depends on the detailed
chemical composition of the peridotite, on thepossible separation of an Fe,S-rich core, and on thetime relation between accretion and melting (seemany papers in Proc. Lunar Sci. Conf.). Whereassome workers have argued for complete earlymelting of the Moon (not necessarily at one time),
others have argued for early melting only of theexterior (e.g. Solomon and Chaiken, I976; Solomon, 1977a ). Considerable cooling occurred in theexterior of the present Moon, and the present
lunar-prete is subject to considerable uncertainty.
The two theoretical geotherms for the outer partof the present Earth (fig. 2, curves ~ and 2) are
generalized for oceanic and shield areas, as
reviewed by Wyllie (197i , pp. 30-2). Pyroxene
thermobarometry of peridotite xenoliths fromkimberlites (Boyd, i973; Nixon and Boyd, I977)
apparently confirms the model geotherms forshield and oceanic areas, but details are controversial. Dynamic models associated with the theoryof plate tectonics indicate complex geographic
variation of the geotherm, as does the existence of'hotspots'. The liquid nature of the outer core andthe solidness of the inner core place constraints on
the temperature, but these are weak because of
uncertainty of chemical composition.
Unfortunately, there are severe problems inestimating planeto-pretes by geophysical calculations principally because of uncertainty in the sizeand position of heat sources and sinks as a functionof time, and in transfer mechanisms. Calculations
for the Earth by Hanks and Anderson (1969), forMars by Anderson (i972) and by Johnston,
McGetchin, and T6ksoz (I974), and for Mercury
by Siegfried and Solomon (I974) , illustrate theproblems with estimating the heat sources. New
developments in the theory of heat transfer during
convection (Daly and Richter, 1978) require
modification of earlier calculations.
Solomon (I977b, I978a, b) reaffirmed suggestions in earlier papers that the timing and extent ofplanetary differentiation is revealed by the presenceor absence ofcompressional or extensional tectonic
features at the surface. Specifically he argues forearly core formation in Mercury, only partial earlymelting of the Moon, and the possibility of late or
sluggish core formation in Mars.
Perhaps simple geochemical arguments will
prove valuable in estimating planeto-pretes. Plainson Mars and Mercury are commonly interpreted aslava flows, and Venus must have been chemicallydifferentiated by crystal liquid processes (see later).
Terrestrial planets must be dominated by various
proportions of O, Si, Fe, Mg, Ca, A1, and S. The
beginning of melting involves complex phase relations, which can be crudely illustrated by reference
to curves in fig. 2. When S and Fe metal are present,
the F e - F e S eutectic might form at c. 1250 ~ from oto c.6o kb with a slow increase of temperature forhigher pressures. The silicate portion of all the
inner planets must be dominated by Mg-richolivine and pyroxene compositions with minor
Ca,Al-rich silicates. Such peridotitic materials
show very complex melting relations in thepresence of H 2 0 , CO2, and other volatiles (e.g.
Wyllie, ~978). Some features are discussed later (p.
22), and it suffices here that 'wet peridotite' tends
to melt between c.I25O and I4OO ~ for mostpressures up to 60 kb, which temperature range isclose to the F e - F e S eutectic temperature (fig. 2).
Anhydrous peridotite could yield basaltic andpicritic liquids at the considerably higher temperatures on the curve labelled 'dry peridotite'. A planetcomposed of 'wet peridotite' could begin to meltnear I25O ~ and most of the volatiles could be
released to the surface taking the radioactive heatsources with them. Subsequent melting of theinterior would require heating to the solidus for'dry peridotite', which process might be difficult toachieve because of paucity of deep heat sources.
The 'wet' differentiation could yield a 'basaltic' or'andesitic' surface, which could undergo complex
crystal-liquid differentiation, ultimately producing
some 'granitic' material. Generalized melting
curves for these materials (fig. 2) lie at considerably
lower temperatures than for peridotite. Subduction
and convection would lead to further complexities
in the planeto-prete.
Melting curves for Fe and enstatite (fig. 2) may
pertain to the differentiation of Mercury, if
Mercury accreted before condensation of feldspar
from the solar nebula. Curve 5 from Siegfried andM I N E R A L O G Y O F THE P L A N E T S

I

2000

I
I
Fe me~tinq

~..5
9"

9/

../i

t" i.I.21-/,...

/'J

9**~dry ,/1500,,;*~"..// .' /.

dry

K

_."

.'" ~ zt
..'"

~

l -

..I

--" "".,~'~ ....-.-- ""~~176

eu~ecl ~cx

IO00l / ~t4/
/ : ..

500) "

200o",,/"we

1 Earth, oceanic
2 Earth, shield
3,4 Moon, Mars
5 Mercury
6,7 Venus, range/basalt
wet
granite

t

2

i

i

4 GPa 6I

8

I0

FIG. 2. Planetopretes and some phase boundaries. Curves [ and 2: mathematical estimates of oceanic andshield geotherms (quoted in Wyllie, i97 D. Curve 5: Mercury model (Siegfried and Solomon, 1974, Fig. 6).
Curves 3 and 4 for Moon and Mars were so close that only one line is given: curve 3: Toks6z and Johnston(i974, Fig. i4, upper bound); curve 4, Johnston, McGetchin, and Toks6z 0974, Fig. 1I, mean A and B).
Curves 6 and 7: qualitative estimates for Venus. Beginning of melting for MgSiO3: Boyd, England, andDavis (1964). Melting of Fe and Fe FeS eutectic: Usselman (x975). Beginning of melting for dry and HzOsaturated basalt: Wyllie (i97 i). Beginning of melting for dry and HzO-saturated granite: Winkler (i 974,
Fig. I8--8). Beginning of melting of dry and wet peridotite--see fig. 8.
Solomon 0974, fig. 6) is approximately consistent
with such an assumption, when account is taken ofcooling of the exterior of Mercury during the past
4 X 109 yr.
The present suggested prete for Mars (fig. 2,
curve 3) is consistent with early melting of 'wetperidotite' and incomplete heating to the curve for'dry peridotite'.
The suggested range of pretes for Venus (fig. 2,
curves 6 and 7) is constrained by the assumptionthat the Venusian interior is similar to that of the
Earth, while the Venusian surface has been held at
c.74o K for the past 4 x I O 9 yr. Perhaps platetectonic processes occur on Venus with complications caused by convection and subduction.

Chemical compositions of the planets are discussed later. Before looking at the processes
involved in the accretion of the planets, it is perhaps
desirable to point out the paucity of direct information on the chemistry and mineralogy of theplanets, and the need to supplement the observations by implications from the chemistry of theplanetary atmospheres. For the Earth direct
information is available only for c.2oo km, which islimited by the peridotite and eclogite xenoliths inkimberlite pipes. All chemical models for deeperseated material ultimately rely on cosmochemical
speculations, and on assumptions about the expulsion of volatile elements, some of which enter theatmosphere. That crystal-liquid differentiation has

J. V. SMITH
occurred on Venus is necessitated by the strong
concentrations of K, U, and Th at Venera landingsites, but equally important mineralogical implications ensue from the presence ofCOz, HzSO4, HC1,
and H F in the Venusian atmosphere. The chemistryof the Martian atmosphere with its strong evidence
for exospheric escape of light molecules has profound implications for the mineralogy and geology,

and provides a warning that the present oxidizing
atmospheres of Earth and Venus may have been
considerably different just after accretion.
The next section attempts to bring together
cosmochemical and cosmophysical concepts usefulin placing limits on planetary compositions andmineral assemblages.

CONDENSATION OF SOLAR NEBULA SET IN CONTEXT OFSTELLAR EVOLUTION AND METEORITES
Who can number the clouds in wisdom? or who can stay the bottles of heaven,
When the dust groweth into hardness, and the clods cleave fast together?
Job xxxviii, 38 9
Although many plausible ideas and calculations
now exist on the presumed origin of the solar
system by condensation of a nebula, Jehovah'schiding of Job is still salutary.
Williams and Cremin (I968) reviewed ideas forthe origin of the solar system, and claimed thatderivation of the Sun and planets from the same gas
cloud leads to the least difficulties in explaining the
dynamical properties, particularly the distributionof angular momentum between the Sun andplanets. Astronomical observations of nebulae,stars, and interstellar dust in the Galaxy provide a
valuable guide to what might have happened in thesolar nebula about 4.6 x lO9 yr ago, as reviewed by

Hartmann (1972) and Strohmeier (I972). Spectrochemical analysis of the Sun's photosphere (Rossand Aller, 1976 ) yielded a bulk composition that fits
well with that of C1 meteorites (except, of course,
for volatile gases). This bulk composition is usedalong with estimates of pressure and temperature inthermochemical models for gas-to-solid condensation (Grossman and Latimer, 1974). The condensed
products are envisaged to coalesce into 'clods', theninto planetesimals, and ultimately into planets bycomplex dynamical processes (Safronov, I972 ),which are complicated by chemical differentiation(fig. 3). The next sub-sections review features
relevant to the mineralogy of the planets andstrongly deflectedby giant planets
growth stunledby Jupiterradioactive and/ few Iorge "lucky" ones IASfERUIL}SI
D/^I,,ICT~.qlM^/r accretion heating,=
I............
I metamorphism,
manydirect capture --~ ]DI A'~-T~ I--I
.
.
~,
me
t
ng
~
'
~,.,,,,-~ I
increasingly I\. . . .
. / I. . . . . .
I I
violenti\\ many,dismtegrotwe /aggregation T\ escaping \capture It"~
I/..-debris . . . . . . . . .
- ~ = 1 ~h~T I telectromagnetic~/~~J
heating?
~
/
f
\
t
I
I gentle/~solar wind removesunbound~ ~],~..ggregation ,/ /~g#orbltIJ' ...... ' J
I N T ~..'~-I ....
~ /Increasingimpact heating~/ J .

o-us-T1alp r [SUPERNOVADEBR,SI....

'rc~176['-~..,~ ~ P O L L O

I:LIST ' ~

/ I
/

_-I I

I

~

~s~lg)~f~cts

n ~J

IPLANETSI 2 ss if)or moons IPLANETS U

ANvil

lAND

Jb~mbardmentfrom J AND

MOONS I c~o~;TLds~

ODIESI'~'~._[ PRES~ ERMoAIIDN-BELT
~s ~

--

\

J

MOONS I

N,

=

,/~

FIG. 3. Summary of possible development of planets and moons from the solar nebula.

M I N E R A L O G Y OF THE PLANETS
meteorites. A general review (Kaula, I975) on 'The
Seven Ages of a Planet' was very useful, as was areview (Kaula, 1977b ) on mechanical differentiation of protolunar material.
Astronomical observations

Cocoon nebulae around young stars are inferred
to contain warm dust from the high ratio of
infrared to visible radiation. Herbig-Haro objects
also have excess infra-red radiation, and show
irregular brightness. Stars passing through the
T-Tauri phase show strong irregular changes ofbrightness and are surrounded by a variable nebulain which particles move up to ioo km/s. The nebula
rotates rapidly and has a strong magnetic field,
which is an important support for theories ofelectromagnetic transfer of angular momentum
from the Sun to the planets. A big mass loss (c.o. t
total mass) from the star is inferred as the solar
wind intermittently floods outwards.
The properties of dust in inter-stellar and intergalactic space were reviewed by Wesson (I974) andin papers collected by Field and Cameron (i973).
Severe technical problems exist in the chemicalinterpretation of light scattered from small dust
particles especially if very small and coated withcondensed molecules (Aannestad, I975; Allen andRobinson, 1975; Day, 1975; Wickramasinghe,
I974; Zaikowski, Knacke, and Porco, I975) , but
the telescopic spectral data are not inconsistent
with the dust consisting of minerals found in
carbonaceous meteorites, expecially phyllosilicates. Many molecules have been detected(Zuckerman, I977) , of which C-bearing ones are
particularly relevant to chemical models forcondensation of the solar nebula. Grain growthfrom inter-stellar gas involves complex processes
that are poorly understood (e.g. Snow, I975), andindeed grains can be destroyed by sputtering andby radioactive and collisional heating near a supernova (e.g. Falk and Scalo, I975).
There are numerous speculations about theorigin of star clusters and about the effects ofsupernovae on the interstellar clouds. Currently,
the old idea of a supernova trigger for the origin ofstars is popular, and indeed some scientists, including Cameron and Truran (i977) , are exploring
whether the gravitational collapse of the solar
nebula was initiated by a shock wave from a nearby supernova. Such a collapsing solar nebula might
contain substantial amounts of 'shrapnel-like'grains and gas from the supernova, and Lattimer etal. (1978) tried to predict the isotopic anomalies
that might survive in refractory grains ejected fromthe various nucleosynthetic zones of the supernova.
Alternatively, the solar nebula might have col-9

lapsed without a supernova trigger, and might haveinherited granular and gaseous debris from anearlier supernova (e.g.D.D. Clayton, I977a, b).
Actually, grains of the appropriate mineral types
occur in meteorites. White inclusions of the Allendemeteorite carry the melilite, spinel, diopside, andperovskite predicted for oxidizing supernova zones
with either C/O < r or (S + Si)/O < i, while enstarite meteorites carry reduced phases including
oldhamite (CaS) predicted for reducing zones withC / O > I or (S--~ Si)/O > I. Observed correlationsof excess 26Mg with 27A1/24Mg in anorthite fromthe Allende meteorite (Lee et al., I977a , b);
Hutcheon et al., I977; Lorin et al., 1977 ) areconsistent with crystallization of anorthite fromaluminium enriched in nucleosynthetic 26A1.
Clayton et al. (I977) interpreted enrichments of 160(relative to terrestrial abundances) in C2, C3, andC4 meteorites in terms of admixture of I to 5 ~oo 1 6 0
ascribed to pre-solar carriers, and Clayton andMayeda (I977) found correlations between O and
Mg isotopic anomalies. Manuel et al. (I972) had
earlier attributed xenon anomalies to nucleosynthesis in a supernova. Recent data on isotopic
anomalies in the early solar system are reviewed by
R. N. Clayton (I978).
Chemical composition of the solar nebula

The simplest approach is to assume that the solar
nebula was spatially homogeneous when condensation began, and indeed there is reasonableagreement between the spectral estimate of thechemical composition of the solar photosphere andthe bulk composition of CI meteorites (Table II)
when account is taken of depletion of volatileelements in CI meteorites. Futhermore, estimateddepletion factors for the Earth, Moon, and presumed parent body of eucrites with respect to C1composition are consistent, within the somewhat
large uncertainties, with the assumptions of a single
reservoir in the solar nebula, and variable depletion
factors related to element volatility and metalsilicate fractionation (Anders, i977; Smith, i977).
However, different ratios of the three isotopes ofoxygen among terrestrial, lunar, and meteoritic
samples (R. N. Clayton and Mayeda, 1978 ) requireat least nine separate reservoirs related by massfractionation and by incorporation of various
amounts of 160. The Earth and Moon haveisotopic ratios consistent with derivation mainly
from enstatite meteorites, but not from ordinary
and carbonaceous chondrites; however, mixturesof several meteorite types are more plausible thanjust a single type. Mechanical admixture of 160
might be explained by incorporation of supernova
debris (preceding section). If ordinary and enstatite

IO

J. V. S M I T H
TABLE I I . Chemical composition of solar photosphere and certain meteorites

H
C
N
O
Na
Mg
A1
Si
P
S
C1
K
Ca
Ti
Cr
Mn
Fe
Ni
Zn
U

a

b

c

d

e

f

g

h

i

atomicatomic

wt ~

wt ~

wt ~

wt ~

wt ~

wt ~

wt

I2
8.62(12)7.94(I5)
8.84(7)
6.28(5)7.6o(15)
6.52(12)7.65(8)5.5o(15)

8.36
7.43(20)
6.28(4o)
8.47
6.36(4)7.61(1)
6.5I(3)7.59(I)
5.57(20)
7.28(7)
5.17(3o)5.15(6)6.44(lO)
4.97(3)5.68(6)5.55(7)7.53(I)
6.25(I3)
4.68(1)-o.4o(12)< 0.040. 5
c.o.o5
36.5
o.35
14.5
1.37
15.7
o.I 3
2.2
o.o3?
c.o.o4
1.7o
o.o87
o.35
o.I 5
25.I
1.4
o.oi
16 ppm

.
.
.
0.4
c.o.2
c.o.o3o.oo 5
27
35
o.82
o.58
lO.7
14.2
o.79
I.OI
16.9
17.o
o.2I
o.I2
5.8
2.0
o.o7?
o.oi?
0.092
0.082
0.84
1.19
o.o57
o.o62
o.32
o.34
0.32
o,22
33.o
27.6
1.8
1.8
0.oo4
o.oo5
II ppmII ppm.
c.o.3
c.o'oo5
39
o.66
15.2
I.iO
18.3
o.II
2.2
o.oi?
o.o9?
1.28
o.o66
o.37
0.24
21.8
1.2
0.006
13 ppm

0.07
o.oo4-c.o.35
23
c.o.5
27
o.oi?
c.o.4
c.4 ppmo.o2?
c.I
c.o.o3c.o.o5
o.14
1.o2
0.02
o.ooi
5 ppm

7.20(15)

5.5(4)
5.16(lO)
6.35(lO)5-o5(I2)
5.7I(I4)
5.42( 16)
7.5o(8)6.28(9)4.45(15)
< 0.6

< 2
< I
3.I
2.2
o.26
c.o.2
46.0
41.5
o.51
o.38
9.56
11.8
0.85
1.o8
IO,5
13.o
o.II
O.lO
5.90
3.4
o.o5?
o.o3?
o.o54
c.o.o5
1.o6
1.34
o.o43
o.o54
o.24
o.31
o.I 9
o.I6
I8.4
21. 9
I.O
1.3
0.030
0.02
9.1 ppm 12 ppm

a. Solar spectrum (Ross and Aller, I976); uncertainty in brackets, e.g. C 8.62_+o.12; normalized to log H = 12; log
atomic, b. CI meteorite; log atomic normalized to log Mg + log Si = 15.2; data sources in Smith (1977, Table 2, col. I).
c. CI meteorites, d. C2 meteorites, e. C3 meteorites, f. enstatite chondrites, mainly E 4. g. ordinary H chondrites, mainly
unequilibrated, h. ordinary L chondrites, mainly unequilibrated, i. enstatite achondrites. Data taken from compilationin Mason (i 971 a). The present table is designed to show important features of meteorite compositions that may be aguide to planetary compositions, and readers are referred to original papers for many complications.
chondrites, along with differentiated meteorites,
are assumed to have formed in the inner solar
system, as is highly probable, then the solar n e b u l a
was n o t isotopically homogeneous. Only if most
meteorites are assumed to be strangers to the solar
system (e.g. the residua of c a p t u r e d comets) can theconvenient a s s u m p t i o n o f a h o m o g e n e o u s primordial n e b u l a be retained.
P e r h a p s the simplest t e m p o r a r y a s s u m p t i o n ist h a t the primordial n e b u l a was sufficiently h o m o geneous in m a j o r elements for c o n d e n s a t i o n
calculations to be applied towards the origin o f theplanets, yet sufficiently i n h o m o g e n e o u s in someisotopes to yield the isotopic variations f o u n d ino r d i n a r y chondrites a n d differentiated meteorites.
The c a r b o n a c e o u s meteorites m i g h t be ascribableto cometary debris. At the o t h e r extreme arepossibilities involving m a j o r inhomogeneities ofthe solar nebula: e.g. M a n u e l a n d Sabu (I975) a n d
S a b u a n d M a n u e l (i976) p r o p o s e d t h a t the solarn e b u l a is c o m p o s e d largely o f s u p e r n o v a debris,t h a t the iron cores o f the inner planets a n d the ironmeteorites derive principally from the centre o f thesupernova, a n d t h a t the outer planets are highly

enriched in elements from the rim o f the supernova.
The p r o b l e m o f the highly reduced enstatite
meteorites is discussed later.
T w o possible causes o f chemical variation in thesolar n e b u l a are m o m e n t u m differentiation, in
which lighter gases tend to m o v e to lower gravitational potential t h a n heavier ones, unless t u r b u l e n t
mixing is strong, a n d effects from the solar wind,which has been jokingly described as a 'magicb r o o m ' allowing theorists to sweep away a n y
u n w a n t e d volatiles !

Physical properties of the solar nebula
U n f o r t u n a t e l y the physical properties c a n n o t becalculated from first principles, a n d all estimates
are controversial.
The original solar n e b u l a must have h a d a mass
greater t h a n t h a t of the present Sun a n d planets.
Assuming t h a t the Sun completed m o s t of its
g r o w t h before the planets accreted, the non-solarpart of the n e b u l a need n o t have been greater t h a n
c.0.01 solar mass if it h a d solar-photosphericc o m p o s i t i o n (Hoyle a n d Wickramasinghe, 1968 ).

MINERALOGY

OF

Safronov (i972) used 0.05 solar mass, but Cameronand Pine (1973) used a massive nebula with I solar
mass. Safronov (I977) argued that such a large
nebula could not have been dispersed by the solar
wind after < o.ot solar mass had entered theplanets, and indeed Handbury and Williams (1976)
have severely questioned the efficacy of the solar
wind in removing even a small nebula. Cameron(pers. comm., I976) has apparently abandoned theconcept of a hot massive nebula.
Ringwood (x975, PP. 563-8) assumed that a small
nebula was relatively cool, perhaps not above
3oo K in the source region of the inner planets andc.2oo K for the giant planets. Such a cool nebula
would allow survival of interstellar grains, and isin complete contrast to the concept of a hot nebula(e.g. Cameron and Pine, 1973) in which initial
temperatures of c.2ooo K resulted in the vapourization of all interstellar grains. Hoyle and Wickramasinghe (1968) discussed the expansion of the solar
nebula by transfer of angular momentum via amagnetic field of io a gauss, and deduced a temperature c.T,(r/R)-~ where ~ is the temperature of the
Sun's surface, R the Sun's radius, and r is distancefrom the Sun. A temperature of I5OO K was
obtained between Venus and Earth and I5O K
between Uranus and Neptune. Lewis's (I974a)suggestion that the chemistry of the planets enforces an adiabatic profile on the solar nebula wascriticized by Kaula (I975), but some kind ofpressure temperature gradient must occur.
A cool nebula would rule out the appealing andhighly developed concept of simple progressive
condensation from c.2ooo K of a chemically uniform gas, as used in many models for chemicaldifferentiation of planets and meteorites. It can beargued that the high-temperature minerals of C3
meteorites formed outside the solar system, Anders(I97 Ia) summarized evidence on the 'condensation
temperature' of ordinary chondrites, the eucrites,
the Earth, and the Moon, but the significance of thevalues of 46o to 6o0 K estimated from the contents
of volatile elements is not clear when one considers
all possible chemical processes in the accretion ofplanets via planetesimals. There is abundantmeteoritic evidence for melting of even small
planetesimals, but this does not necessarily require
accretion from hot condensates. Radiation loss
could allow grains to be cooler by some hundreds
of degrees than the surrounding gas (Arrhenius and
De, I973). Against all these permissive arguments
for cool grains in a solar nebula is the undoubted
evidence for strong depletion of volatiles in the
inner planets (especially for Mercury), and theweaker evidence for enrichment of the Earth inrefractory material (Ganapathy and Anders, I974;
Smith, i977). For the time being it is desirable to

THE

II

PLANETS

keep an open mind on the temperature of the solar
nebula, and especially not to accept uncritically thewell-developed concepts of progressive condensation of a fully gaseous nebula.
Pressure estimates of the solar nebula are also
uncertain, but fortunately are not critical to condensation calculations. The pressure depends onthe degree of flattening of the nebula. Hoyle andWickramasinghe suggested that the density is proportional to r- 3, and Cameron and Pine suggested
a pressure range from io -2 bar at MercUry to I O - 6
bar at Neptune.

Chemical and physical aspects of condensation,accretion, and planetary differentiation
Condensation and aggregation of dust. Althoughsimple progressive condensation of the solar
nebula cannot explain all the properties of meteorites and planets, it provides a useful theoreticalboundary. Grossman and Larimer (1974) summarized the results of many workers in which the earlychemical history of the solar system is developed interms of thermochemical calculations. Solids inthermal equilibrium with the nebula gas condense
when vapour saturation occurs, and change composition or sometimes disappear upon reaction
with the surviving gas at lower temperature. Table
III(a) summarizes the essential features for thermal
TABLE I I I . Effect of C/O on condensation of solar

nebula at io Pa (o.1 millibar)(a) Cosmic o.6
18oo K
16oo-I4oo~4oo-i3oo
70o
5o0
4oo
15~
Ioo-5 o2o. (b) Reduced c.I.O

refractory metals
Ca,AI,Ti oxides
15o0 K
SiC, CaS,TiN,F%C
Fe,Ni,Co and
i4ooFe, Ni,Co
Mg, Si oxides
some Fe --, sulphide I zoo- I IOO Ca,A1,Ti oxides
some Fe ~ oxide
IIoo-Iooo Mg, Si oxides
OH-silicate
water ice
NH3,CH4 clathrates
CH4,Ar ices

Abstracted from Grossman and Larimer (1974) and Larimer(t975).

condensation of a nebula with composition similar
to that in Table II, and with the specific C/O ratioof o.6, and fig. 4 shows some details. The pressure
was assumed to be IO -4 bar, but only small
differences in the condensation sequence occur forpressure changes up to ioo-fold. Prediction ofcondensation of refractory metals near 18oo K andof Ca, Al,Ti-oxides at I6OO-I4OO K fits beautifully
with the observations of tiny nuggets containing

J. V. SMITH

I2

K

1800:
1600
1400

Froclion Condensed
0.4
0.6

0.2

0

' '
l '
Re'
T
, PImelolsREE,, U,Th FAIRLY |iO3r~zAI2SiOz"MqA1204l. . . . . .
FeNiCoICu,___L F'cl'Br'I

400,

,:<::

,

'
~oCO~, Fe304

M(JSO4

I

I

4

}

~

[ VOL~TILES~

SSeTe

j
Cloy minerols

I
I ~,',~,'2~-,i-I q

Ag,Zn,Go~Ge,Sn,Sb

800
600

t

CONDENSATE|

Silicoles

I000

1.0

ICONDENSATIONOF[
SOLAR GAS AT 10"~ ATM~lVIg2SiO4,MgSi03

1200

0.8

L --I

Pb;B,TI In [S Fe304 Fez03

~

-]
[

FIG. 4. Condensation of the elements from a gas of 'solar' composition at Io Pa (io 4 bar)(Ganapathy and Anders, i974). This composition has C/O ratio of o.6. See also Grossman and
Latimer (r974, Fig. a).
Mo, W, and Ru, Os, Ir, Pt (Wark and Lovering,
I976 ) and of melilite-perovskite-spinel aggregates(many references) in the Allende C3 meteorite;
however, Blander and Fuchs (I975) argued for aliquid rather than a gaseous origin of some Ca, A1rich inclusions in the Allende meteorite. Fe, Ni,Cometal, and Mg-silicates (Fe-free forsterite and thenenstatite) should condense under equilibrium at
14oo-13oo K. At higher pressure ( > ~o-3 bar), Ferich metal begins to condense before forsterite, but
most of the metal condenses simultaneously with
Mg, Si-oxides over pressures from IO -5 to io -1
bar. At about 7oo K much of the Fe is removed
from the metal and converted to troilite, FeS, andfrom 75o to 5oo K some of the remaining Fe isoxidized to FeO, which should be incorporated
into the Mg-silicates. Magnetite, Fe304, becomesstable only below 4oo K. Hydroxylated phyllosilicates become stable near 35o K, water ice at
I5o K, various dathrates at Ioo 5o K, and methane
and argon ices near 2o K (Lewis, I972b; Sill and
Wilkening, I978 ). Particularly important are thepredictions that U and Th should be incorporated
into the perovskite, which is a high-temperaturecondensate, whereas Na and K condense as a solidsolution with anorthite only in the mediumtemperature range ~2oo-Iooo K. It must be emphasized that this condensation sequence applies
only to a low-pressure nebula that retains its
bulk composition during cooling, and in which
solid phases maintain thermal equilibrium withthe remaining gas, and do not separate mechanically from the gas.
The simplest attempts (Lewis, I97za; Grossman,1972) to explain the mineralogy of the solar system
implicitly assumed that the chemical compositionof the solar nebula was spatially uniform, that thetemperature and pressure fell off monotonicallyfrom the Sun, that equilibrium condensation
occurred as the nebula cooled, that the solar windterminated condensation by removing remaining
gas, and that the planets accreted only frommaterial in their immediate radial zone without
significant overlap with the radial zones of theother planets. The detailed proposal by Lewis,
which has been reiterated and copied by so manyother investigators that it has almost become adogma, is summarized in fig. 5. Mercury formed at
I44O K and c.Io -2 bar would contain a massive
Fe,Ni core, a small mantle of Fe-free Mg-silicatesand high-temperature Ca,A1,Ti:oxide condensates
but only traces of alkali metals; sulphur and FeO
should be absent unless 'due perhaps to "contamination" by infalling debris'. Venus formed at
940 K and c.io -3 bar would have a relatively
smaller core and larger mantle but each of similar
chemical composition to those in Mercury; sulphur
should be 'virtually absent'. Earth formed at 63o K
would contain considerable sulphur as FeS melt,
while Mars would contain only FeS in its core, andwould contain considerable HzO because it
accreted tremolite. Unfortunately there may be

M I N E R A L O G Y OF THE P L A N E T S
serious problems with this simple model, such as
the presence of water on Earth, the relative densityof Venus and Earth (Ringwood and Anderson,
I976), and especially the strong evidence for anoxidizing CO2-tich atmosphere on Venus. Furthermore, the model does not consider mechanicaldifferentiation, overlap of accretion zones, andchemical effects produced by intense heating of an
accreting or differentiating planet, as emphasizedby Ringwood (I975, Chapter 16). Of course, themodel can be patched, and indeed Lewis (I974b)
proposed that late impacts from comets added
volatile elements to Venus.
Related strongly to the above condensation
model of the planets is the inhomogeneous accretion model proposed by Turekian and Clark (1969)and developed for the Earth by Clark et al. (1972).
They adopted the concept of a hot nebula andassumed rapid cooling and direct accretion of thecondensate into the Earth in not more than lO5 yr.
Effectively the Earth was supposed to accrete its
iron-rich core and Mg,Si, O-rich mantle sequentially from material condensing at successively
lower temperature from the solar nebula. Actually
the first condensates should be rich in Ca,A1,TiI

I

I

I

I

I

I

2000

CaTiO 3x

MgSiO 3

150C

w.
I,-

Fe
1000

?

o_
w
I.-

FeS
500

/
o~c

FeO

ice
0
-7

-6

-5

1
-4

I r
-3

i
-2

1
-I

1
I
0

+1

IOglo PRESSURE (bars)

FIG. 5. Hypothetical condensation model of the planets
based on direct equilibrium condensation of the solar
nebula with C/O = o.6 to a supposed adiabat (Lewis,I972a, Fig. I). See fig. 4 for details of condensation at
IO Pa (= io -4 bar). The symbols represent Mercury(upper), Venus, Earth, Mars, asteroids, Jupiter, Saturn,
Uranus, and Neptune.

[3oxides, and the first Mg-silicates should be Fe-free.
This simple model has been heavily criticized byscientists who favour accretion of planets via
planetesimals, and it will not be considered further
in its simple form; however, heterogeneity ofaccretion probably makes some significant
contribution to the complex mineralogy of theplanets, especially with respect to the outer zone.
Returning to the condensation sequence for
C/O c.o.6 (Table IIIa), at a pressure of IO-4 barhydrogen remains in the vapour until hydroxylated
silicates are formed at about 4oo K. The metal
phases coexist with oxide phases, and there is no
condition for condensation of reduced mineralssuch as Si-bearing metal, sinoite (Si2NzO), andosbornite (TIN). In order to explain such minerals
in enstatite chondrites and achondrites, anothermechanism is needed, probably involving a drastic
composition change of the gas. Larimer (1975)
invoked condensation at low pressure from a solar
nebula with similar composition to that in Table II
except that C/O is near I.O or higher (Table IIIb). If
the enstatite chondrites and achondrites actuallyformed in the solar system, and were not brought inby comets, it will be necessary to find somemechanism for major change of the chemicalcomposition of the nebula; indeed, if the E-asteroids really have surfaces that match enstatiteachondrites, it will be necessary to modify asubstantial region of the nebula. Perhaps someclues may arise from study of carbon-rich stars, forwhich calculations show that reduced species
should occur in the surrounding nebula (Gilman,
1969). This is an extremely important problem thatmust be resolved: in the meantime, it will beassumed that at least the major part of the solar
nebula had solar composition with C/O c.o.6.
Since the equilibrium condensation model reviewed by Grossman and Larimer (1974) relies oncontinuous adjustment between gas and crystalline
condensates, diffusion barriers or mechanicalseparation may cause complications. How would
grains of Fe,Ni become oxidized and the resultant
FeO then enter Fe-free olivine or enstatite? Intervention of melting would ease problems, andindeed there is considerable evidence of melting inplanetesimals and perhaps for dust (occurrence ofchondrules). Blander and Abdel-Gawad (1969)
attempted to explain the origin of meteorites interms of nonequilibrium condensation constrainedby occurrence of supersaturation and solid-state
diffusion barriers, but Grossman and Latimer(I974) argued that meteorite chemistry is moreconsistent with equilibrium condensation.
An alternative approach (Arrhenius and Alfvrn,
i97 I) emphasized ionization of the solar nebula,
thereby ruling out simple condensation based on

14

J. V. SMITH

thermal equilibrium between gas and solid. Hydrogen, O, C1, and Hg would be weakly ionized,
whereas In, T1, Bi, Pb, Pd, Ni, and Fe would bestrongly ionized. Furthermore, solid particles could
have a much different temperature of lattice vibrations than the kinetic temperatures of the electrons,ions, and atoms of the surrounding gas because ofincomplete adjustment between energy transfers byradiative and collisional processes. Ion-sputteringexperiments yielded crystalline and glassy aggregates on substrates at temperatures as low as 300 K.
Several processes of chemical separation in apartially ionized nebula were also considered,especially with respect to a magnetic field associated with a solar wind. Arrhenius and Alfv6n
concluded that the meteorites might range between
high-iron varieties formed from ion-rich plasmaand carbonaceous varieties formed in a vapour
dominated by neutral atoms and molecules.
Meyer (I97I) found that Ca was not fractionatedfrom A1 during sputtering, but that fractionationof Mg and Si with respect to A1 and Ca occurred asa function of the substrate temperature. Electron
diffraction study indicated that enstatite and two
unknown crystalline materials were produced by
sputtering at temperatures even as low as c.4oo K.

Physicalinteractions

involving small bodies.

Because small particles travelling at I 50 m/s innear-vacuum (cAo -3 bar) or air would adhere
spontaneously to targets with rough or dust-coatedsurfaces that 'absorbed' the impact energy, Hartmann (i 978) deduced that no other mechanism wasneeded for accretion of condensed dust into
meteoroid and asteroid-sized bodies. Growthwould be faster for the larger bodies, and would
accelerate when a regolith developed. When thelarger bodies had reached 2 5 km radius, the impact velocities of the swarm would reach 2-5 ~ m/s.
Smaller bodies would become fragmented, andthe fragments would either be swept up by thelarger bodies or driven out of the solar system byradiation pressure. Finally the very largest bodies
would 'run away' from the other large bodies as thecollisional cross-section became augmented by thegravitational attraction.
The greater mechanical strength of iron thansilicates allows the theoretical possibility that
repeated collisions would result in large bodies ofiron separating from a cloud of fine silicate debris(Whipple, 1964; Orowan, i969). Shock metamorphism is a very complex process depending
not only on the relative velocity of the colliding
bodies and the mechanical strength of theirminerals, but also on the relative porosity andinitial temperature of the rocks. On average, however, for rocks starting at c. 30o K, feldspar andsilica minerals begin to fracture upon release fromshock pressure achieved on the Hugoniot curve at
I oo kb and to melt at 2oo kb, whereas pyroxene andolivine are affected only at about twice the pressure(St6ffler, I972, 1974). Kamacite transforms nearI3o kb to e-iron whose distinctive etch patternreveals the solid-state recrystallization (Jain et al.,i972). Shock loading causes hardening, and coherence is retained even at a peak pressure of I Mb,
where kamacite merely undergoes partial recrystallization and shearing offset of inclusions of phosphide and taenite (Wasilewski, I976 ). Perhaps ironfragments could accrete efficiently by welding
together upon impact, so long as the temperature
is high enough for ductile behaviour. The above
pressures could be obtained only for high-impact
velocities (c.I-5 km/s?) when planetesimals had
become large (c.Io2-Io 3 km radius). During earlyaccretion the impact velocities should be too low
for effective mechanical segregation of metal fromsilicate, especially as residual gas would cause
viscous drag; however, at some stage of growthmechanical segregation can be envisaged. If stonyiron planetesimals had differentiated into an ironcore and silicate mantle, impacts would tend tosegregate silicate debris from surviving iron cores.
Magnetic attraction might cause mechanicalseparation of silicate from iron-rich metal (Curie
point: Fe IO43 K, Co I4o4, Ni 631) and perhaps
magnetite. Thus Anders (I97Ib) emphasized that
the Curie points of metal with 5-6 ~o Ni in ordinary and enstatite Chondrites would be somewhat
above 90o K, which temperature is above estimates of final accretion temperatures of the Earthand of parent bodies of meteorites (c.45o 48o K;
Grossman and Larimer, 1974). Particles smaller
than 2o nm would have single magnetic domains,while those larger than I /~m would be multidomained. Only the former should have significant magnetostatic attraction.
Arrhenius and Asunmaa (I973) measured theelectrostatic attraction of lunar soil particles andconcluded that the range of adhesion force (c. IO- 4
to z x Io -3 N) and charge density (IO - 6 toIO lO C/cm 2) would cause particles o.I #macross to adhere upon collision at a relative velocity up to 30 cm/s. Smaller grains could adhere athigher velocities, and pyroxene apparently adheres less easily than feldspar or glass.
Elastic collisions result in less massive bodies
gaining velocity while the more massive bodies losevelocity. The latter should tend to move towards
gravitational highs at the centre of the Sun andplanets while the latter should tend to drift away.
Since metal bodies are stronger than silicate bodies,disintegrative collisions might ultimately haveresulted in Fe-rich metal drifting towards the Sun,and locally towards the planets, while silicate

M I N E R A L O G Y OF THE PLANETS
tended to move away. Perhaps this factor may be
important in yielding a high Fe-content for
Mercury, and a higher Fe content of the Earth thanthe Moon (Kaula, I977a, b).

Physical and chemical differentiation involving
large bodies. Whatever the details, nebular condensates must have aggregated by several processes,
almost certainly at low relative velocity (c.l.5 x
escape velocity of the larger bodies; Hartmann,
I978 ). The resulting 'clods' would sink by gravitational attraction to the medial plane of the nebula,
and would accrete into planetesimals (Goldreich
and Ward, I973; Safronov, i972). Many features of
meteorites are explainable by chemical differentiation in planetesimals at least tens ofkm in diameter,and perhaps hundreds of km; furthermore, themain-belt asteroids range up to IOOO km radius.
Consequently this paper follows the lead of Urey
(I952) in building planets from large planetesimals,
rather than assuming continuous direct accretionof condensed material on to planets (Clark et al.,I972).
The dynamics of two-body interaction are
fundamental to discussion of the physical andchemical differentiation produced by collision andnear-collision. Interaction of two bodies in a single
event leads to either a direct hit, a glancing hit, amiss within the Roche stability limit, or a missoutside the Roche limit (fig. 6).

deflectio..(~ndifferenliol ballistics
for weak body

/
\
X
\

/
~ ~ [2~d,,;~--/'~"~
di ff erentiating/
head-oncottisionplenel~k~,l,..
~/. . . . .

<O) . . . . . . . .

differentiae

captured--.._

planelesimal-I'- - - ' ~~C:"'--i~
.....

~ ~

~ ~
~'-"-debris

10st.~= ringF]G. 6. Cartoon of possible two-body interactions. The
inner zone for head-on collision is enclosed by the twozones for glancing collision and disintegration of a weak
body in the Roche stability zone, and by the region justoutside the Roche zone, for which substantial deflection('sling-shotting') occurs.
The collisional radius Rc of a body with geometrical radius Rg and escape velocity V~for smaller
incoming bodies of non-gravitational relative velocity Viiis augmented by gravitational focusing:R c = R g ( I nI- Ve2/Vi2) 89

I5This effect is shown schematically in fig. 6 by thetapering area marked 'head-on collision'. Reducing
Vi/Veincreases Rc/Rg especially as Viifalls below Ve;
thus for an olivine-rich planetesimal of radiusc.2ooo km with escape velocity c.3 km/sec, thecollisional radius is 2~oo km for Vii at IO km/sec,
28oo km for 3 km/sec, 6o0o km for i km/sec, andextremely large as Vifalls below ioo m/sec. This is amajor cause of the 'tadpole' effect in which thelarger members of a population of planetesimals'eat' the smaller ones (but see later for complications).
Head-on collision could lead to disruption ofboth bodies with dispersal of the debris, but a large
enough body would survive the impact and capture
the smaller body. The products of the impact aregoverned principally by the strength of the bodies,
the impact velocity, and the starting temperature.
Increasing the escape velocity increases the fractionof captured debris; increasing the relative velocity
increases the maximum shock temperature andpressure until ultimately the impacting projectile
explodes into a cloud of vapour, which tends toescape; increasing the ratio of the strength of thesmall body to the large body (which might have
either an uncompacted regolith or a surface weakened by volcanic activity) results in deeper penetration and greater chance of accretion. A thickatmosphere, and especially a ring of debris, could
have major effects on the incoming body and on theimpact products. Modelling major impact events isextremely difficult even when data for natural andman-made explosions are used as a control. Many
pertinent references are listed in Head et al. (x975),
O'Keefe and Ahrens (i975) , and Wetherill (I976).
Recent calculations by O'Keefe and Ahrens (1977b)
suggest that for the present Moon, Mercury, andMars (assumed to have a feldspar-rich surface),
meteorite impacts result in a mass gain only for
impact velocities less than 2o, 35, 45 km/s, respectively, for feldspar-rich meteorites and less than 25,
35, and 4 ~ km/s for iron meteorites. Hence thesesmall planets would retain much less of the material
deflected into the inner solar system by the giant
planets than would Earth and Venus.
Glancing collision formally occurs in an annularregion of width equal to the diameter of theincoming body (fig. 6). The fate of the debris fromglancing or near-glancing collisions was predicted
by Opik (I972), Hartmann and Davis (I975), andespecially by Cameron and Ward (I976) withreference to possible formation of the Earth's
Moon from a Mars-sized planetesimal hitting
at c.Io km/s. The outward-facing part of theplanetesimal might shear off and fragment whereas
the inward part might undergo severe shock including vapourization of the silicate. The combination

16

J. V. SMITH

of differential ballistics with shock-induced dispersal followed by coagulation of only part of thedebris might lead to rapid formation of a refractory, metal-poor Moon near the Roche limit, but it
will not be easy to quantify the qualitative suggestions (see Kaula, I977a, for some scepticalremarks).
The next annular zone involves passage throughthe Roche stability limit for self-gravitation. The
classical mathematical treatment by Darwin (i9o6)
is for a tiny satellite of no mechanical strength
rotation-locked in a circular orbit about the central
body. Inside the Roche radius Rr, the self-gravitation is insufficient to compensate for the disruptive
tidal forces. A simple three-body treatment (e.g.
Haymes, i97i , pp. 207-8 ) gives
Rr = (6Mc/nPs) ~(3)

where Mc is the mass of the central body and Ps is
the satellite density. The two-body tidal treatment
gives
Rr/Rc = 2.45 5(Pc/Ps)~(4)

where Rc and Pc are the geometrical radius anddensity of the central body (Darwin, 19o6). For thepresent Earth (Smith, I974), Rr/Rc ranges from 2.24
for liquid iron (c.7 g/cm 3) to 2.88 for the bulk
density of the Moon (3.34) to near 3 for feldsparrich liquid (c.2.6). Because iron should tend toaccrete earlier than silicate into planetesimals because of mechanical stability, Pc should tend to begreater than Ps: as a rough approximation, theRoche limit is placed at Rr ~ 3Re in fig. 6.
Several complications arise. Mechanical cohesion can prevent gravitational disruption inside theRoche limit (Jeffreys, I947; Mitler, I975); this effectdecreases when the incoming body is hot, andespecially if it is partly molten. Relative rotation of
the incoming and central bodies, and especially anon-circular orbit of the incoming body, increase
the instability region. The last factor is part of thevery complex problem of tidal influence on satellite
orbits (Kaula and Harris, I975, pp. 366 8).
Obviously the Roche limit is merely a theoreticalconcept, a n d many processes can actually occur
within several radii of the central body.
The simplest model for disintegration inside theRoche limit involves passive fragmentation of
the incoming body into pieces that move independently (Wood and Mitler, I974) with thosenearest to the central body going into ellipticalorbits about it, those farthest from the body being
lost into hyperbolic orbits, and those in an intermediate zone going into such elongated orbits thatescape is likely. For capture of a significant fraction, V~ must be less than 2 km/s, which for the
Earth requires that the semi-major axis of an

incoming body must lie between 0.8 and 1.2 AU(Kaula and Harris, I975). Unfortunately, thesimple calculation is only a mathematical boundary condition because fragmentation would not
be passive, the fragments would not simply assume
the velocity of the parent planetesimal, and much of
the material that flew off into heliocentric orbit
would ultimately impact the Earth and any
surrounding cloud or moon(s) on succeeding
orbits.
The fourth zone in fig. 6, which must overlap the
Roche zone for strong bodies, involves merely
orbital deflection and trivial tidal heating. In a
simple two-body encounter the effect is apparently
trivial but taken in the context of the entire systemof bodies orbiting the Sun this 'slingshot' effect isextremely important as it changes the eccentricityof the orbit thereby causing overlap of accretion
zones of the planets (see later).
Finally the simple two-body encounter must be
extended to include collision with any debris ring(including possible moons) in orbit around a growing planet or planetesimal. Kaula and Harris (~973)
showed that collision of a Moon-sized object with
2 x IO- 5 of its own mass in Earth orbit dissipated asmuch energy as by the tidal effect during a noncollisional near-encounter. Ruskol (I972) extended
earlier papers on the origin of the Moon by
discussing how an incoming body could either
destroy bodies in geocentric orbit or be captured toaugment the orbiting material. In general the larger
the body or bodies in Earth orbit, the greater theirchance of survival and the less the likelihood of
impact debris ending up on Earth. O f course, a very
thick, deep atmosphere such as the present atmosphere of Venus, and particularly the early planetary atmosphere postulated in the single-stage
model of Ringwood (I966b), would contribute to
the effect of a debris ring.
Attention became focused on disintegrative capture because of the need to explain the coupling of
the low iron and volatile contents with the somewhat high content of anorthite inferred for the bulk
Moon with respect to the Earth. Numerous authors
listed the possible chemical effects that could
accompany the dynamic effects of disintegrative
capture (e.g. Kaula, i97I, I977a; Ruskol, t977) , but
quantification is difficult. An incoming body might
have chemically differentiated into a metal core anda silicate-rich mantle and crust, thereby permitting
several types of mechanical separation such as
stripping off the outer part by collision with debris
in Earth orbit (Kaula and Harris, I973), ballistic
differentiation (Wood and Mitler, I974), andmomentum differentiation of massive Fe-rich fragments with less massive silicate-rich ones (Smith,I974; Kaula, ~977a). Major loss of volatiles couldM I N E R A L O G Y OF THE P L A N E T S
occur in any collision involving relative velocities
greater than a few km/s: of course, the greater thetemperature rise the greater the range of volatileelements that would be significantly depleted.
Although these processes cannot be quantified itwill be assumed that they typically occurred during
the interaction of planetesimals with growing
planets and that they substantially modified theprimary condensates from the solar nebula.
Before facing up to multi-body problems it isconvenient to mention briefly the problems of tidal
interactions. Once a moon has formed it causes
tidal perturbations on its planet, and undergoes
perturbations itself. Even ignoring impacts by lateaccreting bodies the tidal evolution can be complicated (Kaula, 197I; Kaula and Harris, 1975). A
satellite orbiting sufficiently slowly with respect tothe rotation of the central body will recede if theangular momentum vectors are parallel (i.e. prograde rotation of the satellite), and will spiralinwards for retrograde rotation. If the central bodyis rotating very slowly with respect to the orbital
period of the satellite, the latter can be captured
whatever its direction of rotation (Burns, 1973;
Counselman, 1973; Ward and Reid, 1973), especially when the central body is close to the Sun.
Indeed Singer (197o) postulated that the slow
retrograde rotation of Venus results from captureof a retrograde moon. Tidal recession of a satellite
results in rotation-locking such that the satellite
always shows the same face to the central body,
except for small librations. Many complications
arise if the two bodies change mass and velocity
during simultaneous accretion (Harris and Kaula,1975).
Because of the existence of mechanisms for
destruction of satellites, the absence of satellites
around Venus and Mercury, and the small size ofthe ones around Mars, will not be taken to imply
that substantial debris clouds and perhaps large
moons never existed around these terrestrial
planets. Furthermore, the presence of only one
Moon around the Earth does not preclude theearlier existence of several moons and a debris
cloud. Indeed Ruskol (i96o , i963, 1977) and Opik(1972) developed detailed mechanisms for generation of the Earth's moon by aggregation of several
moonlets produced from a thick debris cloud, andHarris (1978) presented further considerations.

Planetary accretion: dynamics, time scale, andheat sources. The dynamical processes of planetary
accretion from planetesimals are too complex to bereviewed satisfactorily here, and indeed the calculations are highly dependent on the uncertain physics
of inelastic collisions. Readers are referred initiallyto the book by Safronov (1972), which covers theideas developed in the I96OS by Soviet mathemati-

17clans, and to the book by Opik (i976). The present
account briefly abstracts the summary in Safronov(1977) , and refers to some recent amendments.
At the beginning of accretion of the planetesimals the collision probabilities can be calculated bystatistical models. As the protoplanets grow thegravitational potential becomes non-uniform, andthe two-body collision outcomes become more
complex, thereby resulting in the mathematicalmodels becoming less accurate. Ultimately most of
the mass has entered the planets, and the motionsof the surviving planetesimals can then be modelled(again statistically) by Monte Carlo calculations,whose accuracy depends on the validity of theassumed mass-space-velocity distribution of theplanetesimals and on the rate of accumulation ofsmall errors caused by approximations in the
successive orbits.
A basic equation for capture of small planetesimals by head-on collision with a protoplanet is:

dm/dt = 4rce2(I + Ve2/]/2).G/P
where dm/dt is the increase of protoplanetary mass
with time, Rg is the geometrical radius, ~ is theescape velocity, g is the mean relative velocity ofsmall planetesimals with respect to the planet, andcr is the surface density of planetesimals in the zone
swept out by the protoplanet as it orbits around the
Sun with period P. Of key importance is theincrease of capture cross-section with escape velocity as described in the preceding section, leading toa few large bodies capturing the smaller ones. If theplanetesimals have exactly circular orbits in theecliptic plane each protoplanet could sweep upmaterial only from a torus of area nR2(1 + ge2/g2),
and from a larger volume for bodies precessing
outside the ecliptic plane. Thus continued
growth depends on capture of material with noncircular orbits, and here is the crucial and probably
insoluble problem. Elastic collisions between
numerous small bodies cause 'jostling' and a tendency for the bodies to adopt circular orbits withvery small relative velocity (as in Saturn's rings).
Accretion of planetesimals may begin by gentle
aggregation into an enormous number of small
bodies ( < I km radius) composed of lumps thathave fallen into the plane of the ecliptic after
condensation. By chance a few of these small bodies
will coalesce thereby reducing the population, butplanets can form only as the planetesimal orbits
become non-circular, and indeed the final stage ofaccretion must involve capture of planetesimals
with orbits at least as eccentric as half the distanceto the next planet.
Unfortunately it is not possible to accurately
calculate the changes of P and orbital eccentricity
as the planet grows. The key factor is the relative

I8

J. V. S M I T H

extent o f inelastic and elastic collisions in two-bodyinteractions. Deflection from near-collisions results
in increase of V and orbital eccentricity of surviving
planetesimals,
whereaseach
disintegrative
encounter results in removal of a planetesimal withno effect on surviving planetesimals. Only thelarger bodies can cause significant deflection, andindeed the change of ~ is governed almost entirely
by interaction with the largest body in each 'zone'.
The largest body will retain a near-circular orbit
because of statistical averaging of captures.
Safronov provided a mathematical framework
for calculating the temporal variation of size distribution of protoplanets and planetesimals ineach planetary 'zone'. Qualitatively, the growth ofthe planets is envisaged in fig. 7. Planetesimals growin each planetary 'zone', but the single protoplanet
in each zone grows faster than its competitors. It
begins by capturing near-by competitors with low
relative velocity and ends by capturing survivors
with high relative velocity. The 'feeding zone' of
each planet flares out, and if V increases sufficiently
the feeding zones overlap, and the planets 'compete'
for the survivors. Safronov approximated ~ 2 =V~z/20 where V~ is the escape velocity of the

-,-

Now

protoplanet in each feeding zone. Safronov estimated that 0 rose to c.3-5 in the Earth zone giving
final accretion from planetesimals with V c. 4 km/s.
That Mars is a surprisingly small planet compared to the expected mass in its feeding zone, andthat a planet did not accrete in the asteroid belt, are
two obvious consequences from increase of P by
gravitational interaction with near-by Jupiter.
Indeed most of the mass in the Mars and asteroid
zones might have been ejected from the solar
system or deflected into other planetary feedingzones (Shoemaker, 1977b; Wetherill, i977) , but noconvincing calculations have been made. Perhaps
Mercury was also stunted by effects related to theproximity of the Sun, whose gravitational pullwould result in high velocity of material perturbed
into the Mercury zone.
The extent of overlap of planetary feeding zones
is a vexing question, as is the mass-orbital distribution of surviving planetesimals responsible for thelate bombardment of the inner planets. Wetherill(1976, 1977) made M o n t e Carlo calculations of theplanetesimals surviving after accretion of 98 ~o ofVenus and Earth, F o r s
km/s the swarm was
captured with a half-life ofc. 107 yr at - 4.5 x I o 9 yr,

comets---~debris-4.0/ , f / ". ,/ 1

/ Y

\\~\

Gyr -4.5

9

-

I

/

_%,-x.\I

/

/

~

\

,\

.

/" .//

late--

9/ ~ ' >

-

,,'/

n.

.

.'~

I / ~
v

1

-

bombordment~
.

.

.

.

.

differenTiaTea\ ~X~M#,__~aste r~ ? //-4.65 r
S

/

I
I

/-

, f

,,yVV,,.__'~,_/___~7 - comets?vVvvwl Y Yvvvv.W vv,v V

Me V E Mo

A(distance)112
FIG. 7. Cartoon of growth of planets from planetesimals. Of the many planetesimals only a few survive to
become the planets and asteroids displayed at the square root of distance from the Sun. The feeding zones
widen with time as V, the mean encounter velocity, increases, and planetesimals are captured by the growing
planets. Feeding zones ultimately overlap and most of the accretion has occurred by -4.5 Gyr for the inner
planets. Jupiter (and the other giant planets, not depicted) ejects planetesimals both outwards to become
comets, and inwards (not shown) to bombard the inner planets and especially the asteroids and Mars. The
late bombardment of the inner solar system by the near-exponential decaying flux of planetesimals from theouter solar system is depicted schematically. The present flux of comets and asteroidal debris is much lower inmass than early fluxes. Possible disintegrative capture of a large planetesimal by the Earth to form the Moon,and possible transfer of planetesimals from the inner solar system to the asteroid belt to form 'differentiated'
asteroids are highly speculative.

MINERALOGY OF THE PLANETS
c. 1 0 8 at - 4.o, and c. 1 0 9 at - i.o. Calculations of theresulting mass distribution yielded a mean lifetime
ofc.3 x lO7 yr for IO24 g bodies and c.2 x lO6 yr for
IO22 g. Using Wetherill's procedure, Hartmann(I976) estimated that about half of the late accretion of Mercury, Venus, Earth, and Mars camefrom planetesimals that began in the feeding zones
of other planets. Furthermore, he suggested that
the Moon received 6o ~ of its late accreted massfrom planetesimals that grew in the Venus zone.
These calculations refer only to an assumed
population of late planetesimals, and only further
study can test whether the feeding zones of theplanets would overlap significantly at an earlier
stage of growth. In this lecture it will be assumed
that at least 9o ~o and probably at least 95 ~ of themass of an inner planet is derived from material
accreted near its present orbit, and that the bulk
composition is indicative of the condensed materialof the local zone of the nebula: however, it would be
wise to investigate the consequences of greater
mixing.
Safronov (I977) deduced that the Earth accretedwith a half-life of c. I O 7 yr and that 98 ~o accretionoccurred in c.Io 8 yr. Interpretation of radiometricchronology of the early solar system is not easy,
and may need revision if the solar nebula wasincorporating supernova debris. Wetherill (1975)
interpreted the evidence for meteorites in terms of
rapid accretion of small bodies into planetesimals(C.lOO km radius) in c.Io 4 yr with subsequentheating and cooling o v e r l O 7 - I O 8 yr. Roughly 10 8
yr is needed from the nebular period to establishment of the reservoirs of lead in the Earth.
A serious problem for simple accretion models is
the relative time scale for accretion of the inner andouter planets. The existing accretion models of
Safronov-type lead to much longer accretion times
for the outer than for the inner planets, as can be
deduced simply but not rigorously from the a/P
term in the earlier equation for dm/dt. If the Earthaccreted in io s yr, Mars and the outer giant planets
would have required at least lO9 yr, according to
the simple Safronov model (Safronov, 1977;
Weidenschilling, 1976). Since accretion of the outerplanets should have resulted, in deflection of many
planetesimals into the inner solar system, and thesimple interpretation of the morphology of Marsassumes essential completion of accretion in time
for a differentiated crust to retain a record of thesame population of bodies as formed the lunar
basins, it is desirable to investigate models for more
rapid accretion of the outer planets, and perhaps
for the inner planets as well. Weidenschilling
pointed out that the Safronov model actually givesa low estimate of accretion time because it neglects
the difficulty of capturing planetesimals at the edge

I9of a feeding zone, thereby accentuating the
problem.
The Safronov model predicts that the mass
distribution of bodies in each feeding zone is verystrongly biased. The protoplanet is by far thelargest body, and the simple model suggests thatthe next largest body has not more than c. I ~ of themass of the protoplanet. If true, this would rule out
the model of disintegrative capture of a Mars-size
body from whose chemically differentiated debris
the Moon was formed. However, Hartmann andDavis (I975) made further calculations, whichsuggest that the second-largest body in the Earthzone might have a radius of 50o to 3ooo km, which
upper limit corresponds to a volume five times thatof the Moon. Wetherill (1976) questioned theaccuracy of the Safronov-type model and suggested change in the parameter 0. I suggest thatlarge bodies might develop at the interfaces ofplanetary feeding zones because of lack of competition with protoplanets when V is low, and recommend investigation of the possible capture by the
Earth of a large body growing between Earth andMars (fig. 7).
The present masses of Mars and the asteroids aremuch lower than the estimated mass of non-volatilematerial in the relevant zones of the solar nebula(cf. Wetherill, 1976; Weidenschilling, 1976) andprobably an enormous mass of material was ejectedfrom these zones by Jupiter (Ip, 1978). Most of themass would be ejected outwards with some ending
up as comets and some colliding with planetesimals
in the outer solar system to end up in the giant
planets, but a significant fraction should haveended up in the terrestrial planets. Together withthe material deflected by Saturn, Neptune, andUranus, there is scope for developing theories ofbombardment of the terrestrial planets by a considerable number of volatile-rich bodies eventhough most of the material would arrive at suchhigh velocity (especially at Mercury) that perhaps
only the larger planets Earth and Venus would be
able to retain the collision debris while the smaller
planets Mercury, Moon, and, to a lesser degree,Mars, would lose most of the debris because of theirlower ratio of escape velocity to impact velocity.
Consequently it is not unreasonable to invoke thequalitative possibility of late heterogeneous accretion of volatile-rich material on the larger terrestrial planets. Furthermore, it is possible to envisagethat the planetesimals of the inner part of the solar
system (Mercury-Mars) were almost free of volatiles (as indicated by many meteorites), and that
substantial amounts of volatiles condensed only
from the asteroid zone (or perhaps Mars zone)
outwards [this refers to volatiles condensing below
600 K, and perhaps as high as 1IOO K]. Indeed Sill

J. V. SMITH

20

and Wilkening (i978) concluded that the atmospheres of the terrestrial planets cannot be explained by condensation models in which the Earthaccreted at 5oo K, and suggested that the noble
gases of the Earth were provided by about ] ppm ofclathrates, which originally condensed in the outersolar system.
Finally, heat sources need discussion (Safronov,~977; Weidenschilling, i976 ). Conversion ofgravitational energy of infalling bodies into heat is
trivial for small planets and planetesimals, andbecomes important only when planets have grownlarge. Estimates of V from Safronov-type models
are too low for significant heating from the collision of planetesimals whose velocities have notbeen 'pumped up' by mature planets or the Sun, but
can provide substantial heating when ~" reaches
several km/s. The higher values of V envisaged by
Wetherill (i 976) may be significant. Heat loss froma planet's surface becomes more important whenaccretion slows down. Models of radioactive heating do not lead to early melting ( a t - 4 . 5 • Io9 yr) ofa terrestrial planet unless the accreting planetesimals are hot ( > Iooo K), whereas it is easy to findconditions for late melting of at least the outerzone. Actually there is so much mineralogicalevidence for high temperatures in differentiatedmeteorites that even small planetesimals (c.[o z to
1o 3 km radius) of the inner solar system must havereached high-metamorphic temperatures or evenmelted. Perhaps short-lived radioactive elements(e.g. 26A1) were responsible, as proposed by many
scientists (see Herndon and Herndon, I977), or
perhaps some other mechanism or mechanisms(including electromagnetic induction heating;
Sonett et al., 197o ) were responsible. For the rest of
this lecture, it will be assumed that most planetesimals in the region of the terrestrial planets reachedor nearly reached melting temperatures, but notnecessarily those in the region of the giant planets,
and that terrestrial planets accreted from hot
material and underwent an early melting with outsweeping of the elements that entered H20-CO2bearing melts before accretion had ended.
Chemical differentiation inside planetesimals andplanets. This section begins with a review of phaseequilibria particularly pertinent to crystal-liquid
differentiation in terrestrial planets and meteorite
parent-bodies.
Figs. z, 8, and 9 show the important phaseboundaries for the major minerals and rocks expected in terrestrial planets. Crystal-liquid differentiation of a bulk composition composed only of
enstatite (MgSiO3) and iron metal would begoverned by the melting curves (fig. 2) for iron
(Sterrett et al., I965) and enstatite (Boyd et al.,I974). Substitution of IO% Ni and o.5 wt% Co
would reduce the melting point of metallic iron by

kb
2000

20
I

40

60

. ~"//h I

~//i

80

I

kb

I00

I

20

!0001

Ic,c~,~

I~ ~,~k~r;~>r~'lJQzI.,~',Cc

40

I.~ - [I_( R

O [~/~
I
o

80

100

I

teirehedrol C

,

/a.e

f

/
//.
j

60

phlogopitestability,,)1
IOOI "'2
/
K
"3

/

(

"~..~.
excess water]'~--amphibole...........stability/'

1000 ~l'~-C"c~#growth ofzeolites/

1 Tr+Fo~
En+Di+H20
2 Do+Ta+4Qz~Tr + 4 COg
3 Se~Fo+Ta+H20

Iting ofwater ice
I

I

I

4

6

8

GPa

h

o, -

peridotite with 0.1% CO2
I Fo+ Opx +Cpx+Do-*'- L
20px + Do~ Fo + Cpx*C023 Sp+ Ra--~-La+COz

4 3Do+Di-b'2Fo+4Cc*2CC~__
500
9Fe(CO)s condensation~COz&CO melting

~o

oI

0

[

2

I

5 Tr + IIDo--)-8 Fo +13Cc + 9C02+ HzO
6 Ta+2Do +4Qz--~
If +4C02

II

4 G Pa 6

8

I0

FIGS.8 and 9: FIG. 8 (left). Possible phase relations involving H20. Di diopside, Do dolomite, En enstatite, Fo forsterite,
Qz quartz, Se serpentine, Ta talc, Tr tremolite. F~G.9 (right). Possible phase relations involving CO2. An anorthite, Cc
calcite, Cpx clinopyroxene, Di diopside, Do dolomite, Fo forsterite, Gr grossular, Ky kyanite, La larnite, Me meionitescapolite, Opx orthopyroxene, Qz quartz, Ra rankinite, Sp spurrite, Ta talc, Tr tremolite.

M I N E R A L O G Y OF THE P L A N E T S
only IO K. Forsterite (Mg2SiO4) has an even higher
melting point than enstatite. Thermochemical
models of the solar nebula predict condensation of
Ca,A1,Ti,O-rich mineral assemblages prior to condensation of Mg-silicates and Fe,Ni metal.
Detailed melting studies have not been made on thespecific bulk composition implied for Mercury by
the model of Lewis (1972a), but the phase equilibria
for 'peridotitic' compositions have been studied
extensively. Peridotite compositions (Wyllie, 197I )free of H 2 0 and CO z begin to melt along the curveafgh (fig. 8) producing increasing amounts of melt
as the temperature rises. The melt is 'basaltic' incomposition near the solidus and becomes increasingly richer in pyroxene and olivine components as
melting proceeds. Below the solidus the peridotitic
compositions crystallize into assemblages ofolivine and pyroxenes with either plagioclase (lowpressure field: pl-pe outlined by aft), or spinel(medium pressure field: sp-pe outlined by ifgj) or
garnet (high pressure field: ga-pe outlined by jgh).
At even higher pressures peridotitic compositions
yield oxide and perovskite minerals. Whatever thedetails of the condensation of the solar nebula andaccretion of planetesimals, some approximation to'peridotitic' composition can be expected for all the
inner planets, and the 'dry peridotite solvus' marks
an upper limit to the beginning of melting. See
the Proceedings of the Lunar Science Conference
for experimental details of many specific compositions.
The phase relations change drastically and become very complex indeed as soon as Na, K, HzO,
and CO2 enter the planetary composition. Sodium
can be accommodated in plagioclase and pyroxene
minerals, but will tend to enter magma produced bymelting of peridotitic composition. Potassium does
not enter the dense minerals composing dry peridorite (except at pressure over 40 kb) and is strongly
concentrated into magma or slightly into plagioclase. Addition of H 2 0 and CO2 causes drastic
effects whose details have been explored only in thelast decade. The present account is based on aforthcoming review of the peridotite-COz-H20
system by Wyllie (i978) and utilizes a series ofstudies including Eggler (I976 , i977) , Wyllie andHuang (1976a, b; I977) , and Mysen (I977).
In the presence of H20, Na, and K (as for
peridotites on Earth, probably on Mars, andpossibly on Venus), amphibole and mica become
stable, as delineated by Boettcher and co-workers.
The stability limits of these two minerals vary withmany factors such as Pn~o, Mg/Fe, Ca/Na, etc., but
the two generalized curves in fig. 8 are within IOOK
and 5 kb of most conditions to be expected in Earthand Venus. Mg-rich mica (phlogopite) is stable tohigher pressures and temperatures than amphibole,

21

and a K-bearing peridotite will begin to melt along
the curve abcde (fig. 8). Roughly speaking, H 2 0 canbe retained by mica and amphibole in a wet planetu p t o C. I ZOO-1400 K. Titanoclinohumite and titanian chondrodite are very rare minerals in the Buell
Park kimberlite, and might be important in subduction zones of the upper mantle (Akimoto et al.,I977).
Presence of CO 2 leads to phase relations that areeven more complex, but the main features areillustrated in fig. 9- Melting in the presence of o. I ~o
CO2 begins along the curve abdegh, whose temperature varies only about lOO K for the pressure
range o to 4o kb. The maximum at d was notproperly recognized in early work, and is importantfor models of explosive loss of CO2 from rising
peridotitic melt. Fields abc, cbdef, and fegh outline
the stability of plagioclase-, spinel-, and garnetperidotites. Dolomite-rich carbonate (abbreviatedDo) is stable in the region bounded by curve gh (no.
I) for the melting of forsterite, two pyroxenes, andDo, and by curve gi (no. 2) for the transformation oforthopyroxene+Do to forsterite, clinopyroxene,
and CO2. Curve 2 extrapolates into the cluster ofcurves 4, 5, and 6 for the breakdown of dolomite in
the presence of various silicates at low pressure (seelater). Presence of both H 2 0 and CO2 leads tosubtle complexities reviewed by Wyllie (I978).
Sulphur is a common constituent of meteorites
as troilite, pyrrhotine, or sulphat e minerals. If it
occurs as troilite in accreting planets, the Fe,Ni
metal could react with the troilite at the Fe,FeS
eutectic (Usselman, I975), which changes composition from z7 wt ~ S at 3o kb to 24 ~o at IOO kb. The
Fe,FeS eutectic stays at c. 125o K for pressures up to
55 kb, and increases slowly for higher pressures.
Limits can now be placed on the initial melting ofa planetesimal or planet containing volatile elements. I f S is present arid K, Na, H 2 0 , and CO2 areabsent, melting could begin at c.I25o K leading tosinking of a heavy Fe, S-rich liquid (together withassociated Fe,Ni metal) to form a core surrounded
by a mantle of unmelted peridotite. Further heating
would lead to 'peridotite' melting near the 'dry
peridotite solvus'. Addition of N a and K to theperidotite would lead to somewhat lower melting
temperatures, but not as low as the Fe, FeS eutectic,
while addition of COz and especially of water
would lead to considerably lower temperatures.
Indeed hydrous K-bearing peridotite would meltnear the Fe,FeS eutectic for pressures of Io to
3o kb, and 'basaltic' melt could rise as Fe, S melt
descended. Actually there are minor complexities
involving the mutual solubility between basaltic
and sulphidic melts. Perhaps such a double differentiation occurred as the proto-Earth was growing,
but perhaps water was absent in the early stages

22

J . V . SMITH

with differentiation involving mainly the sulphidic
component. For the double-differentiation,sulphur- and metal-loving elements would concentrate into the core while magmaphile elements
would enter the 'basaltic' fraction. Particularly
important would be the fate of radioactive elements: for double-differentiation, the U, Th, and K
would probably concentrate almost entirely into
the 'basaltic' fraction, but for H20-free differentiation, these elements might be fractionated substantially into the sulphidic liquid, especially if a highly
reduced state leads to an increased chalcophilic
behaviour of these elements (Smith, 1978).
The pressure-temperature curve for a protoplanet depends critically on the temperature ofaccreting material, the collisional heating of accretion, and the heat loss from the surface. All threefactors are highly uncertain, but most simple
models yield a temperature maximum closer to thesurface than the centre, often about one-tenth to
one-quarter down. Sinking Fe,S-rich diapirs andrising 'basaltic' diapirs can be envisaged as partial
melting begins. Gravitational energy is released bydiapiric separation, and catastrophic heating anddifferentiation might occur for an Earth-sizedplanet, as suggested by Urey (i 952) and emphasizedby Ringwood (I966b, 1975). However, such catastrophic differentiation need not occur for a protoplanet that begins to differentiate early in its
growth.
Accretion of large planetesimals (c. I o 2 to I o 3 km
radius) should cause catastrophic crystal-liquid
differentiation in large regions of a growing planet,especially if the planet is at near-melting temperatures. Large overturns might occur, and the possibility that the outer part of the growing Earth wasessentially 'stirred' by impacting bodies should beconsidered.
A 'basaltic' veneer on the Earth could undergo
many complex differentiations as secondary processes (igneous, metamorphic, and ultimately sedimentary) occurred. Two important considerations
are the relative melting relations of wet vs. dry'basalt' and of wet vs. dry 'granite', for which
generalized curves (fig. 2) are abstracted from
Wyllie (I97I, I978), Stern and Wyllie (I973) , andWinkler (1974). These curves are not pertinent to
the primary differentiation of planets because thestarting compositions are dominated by olivine andpyroxene, which components must be segregatedinto a mantle before the feldspar minerals andquartz can dominate the phase relations of upward
segregates.
After cessation of bombardment by planetesimals the Earth apparently developed a 'granitic'
crust and water-borne sediments b y - 3 . 8 x IO9 yr.
The development of plate-tectonic processes iscontroversial, but present evidence indicatessubduction to at least 5oo km and perhaps deeper(Davies, 1977). Theoretically, one can imagine thata cooling planet could resorb its hydrosphere (andsome of the atmosphere) when the radioactive
heating diminished enough to allow hydrous
minerals and carbonates to be stable at considerable depth. For this to happen subduction mustcarry volatile-rich sediments deep into the planet.
Much of the lower crust and upper mantle of thepresent Earth would adsorb H 2 0 and CO2 iftransported downwards from the hydrosphere andupper crust. Such readsorption could occur even if
parts of the upper mantle were still outgassing
volatiles by explosive volcanism. If magmas fromthe upper mantle could react with a granulitic lower
crust during transport to the surface, they could
lose their volatiles: indeed the concept of volatiletraps corresponding to the stability limits of mica,
amphibole, and other minerals would become
viable.
Much further work is needed to explore thepossibilities of primary and secondary crystalliquid differentiation of a growing terrestrial
planet, and the even more complex possibilities ofreturn of volatile elements into a dying planet.
Further phase equilibria are now explored to place
limits on mineral reservoirs for H 2 0 and CO2.
Fig. 8 pertains to H20. The melting temperatureof H20-ice varies slightly with pressure (Clark,1966) and may provide an effective lower boundary
for the crystallization of zeolites. On Earth large
beds of zeolites formed from alteration of volcanicash (Munson and Sheppard, i974). The stippled
area in fig. 8 for growth of zeolites under metamorphic conditions is limited by the reactions ofbreakdown of heulandite, laumontite, and wairakite (Winkler, 1974, Fig. I2 ! I). Analcime is stable
only to about 480 K in the presence of quartz, but isstable to magmatic temperatures in some alkaline
compositions (Kim and Burley, I97I ). The breakdown of serpentine to forsterite + talc + water(curve 3) is one of many possible reactions involving hydroxylated layer silicates (Winkler, i974, Fig.
II--5). For amphiboles the stability depends
strongly on the chemical composition. Tremolite
and forsterite react to enstatite + diopside + H 2 0some 3oo K higher than the breakdown of serpentine, and anthophyllite breaks down to enstatite+ quartz + H 2 0 at a similar temperature to that oftremolite (Winkler, I974, Fig. I I--7). Aluminiumrich hornblendes from peridotitic rocks are stableto even higher temperatures, and the dashed curve
marks the approximate limit based on many experimental syntheses (Mysen and Boettcher, i975). For
micas the maximum temperature stability is for Fefree phlogopite, which breaks down to forsteriteM I N E R A L O G Y OF THE PLANETS+ liquid in the presence of enstatite (Modreski and
Boettcher, 1972 ). Fe-bearing phlogopite found in
xenoliths from the Earth's mantle is stable to aslightly lower temperature. If the Martian rocks arevery Fe-rich the stability fields of both mica andamphibole would be lowered several hundreds ofdegrees with respect to curves in fig. 2. Phase
relations for muscovite (and talc) are given by
Wyllie (1973) and Huang and Wyllie (1974).
All the above phase equilibria are for PH2o =Ptotal. Phase relations for systems containing both
CO2 and H 2 0 are very complex, as reviewed
in Winkler (~974), but the reaction dolomite + talc
+ quartz going to tremolite + CO2 (fig. 2, curve 2) isfairly straightforward because it applies only to ahigh ratio of CO2 to H20.
Many other minerals should be added to fig. 8
including the epidote and chlorite groups, montmorillonite, and staurolite. Systematic perusal of
Winkler (1974) will reveal details which cannot begiven here. Comparison of figs. 2 and 3 shows thattheoretically many minerals could store water in all
the terrestrial planets except for Venus in which
only the amphibole and mica groups are serious
candidates. Of course, water-bearing minerals
would not occur if the planet were bone-dry, asapparently is the Moon.
Because of the presence of CO2 in Mars, Earth,and Venus, and the high abundance of C in the Sun,a skeletal review of the phase relations of C-bearing
species is useful (fig. 9). Carbon monoxide and CO2
melt near 74 and 216 K at low pressure; the former
boils at 82 K and the latter sublimes at I95 K(Weast, 1969). The melting point of CO2 increases
to 4oo K at 15 kb (Clark, 1966, p. 373)- Methanemelts at 91 K and boils at lO9 K (Weast, 1969),
which values are important for the giant planets.
The CO1.6H20 clathrate, isostructural with craigite, breaks down to H20-ice and CO2 at I5O K ando.oo4 bar and at I9O K and 2 bars (Miller andSmythe, I977).
Iron occurs as a vapour deposit in lunar rocks,
and transportation as pentacarbonyl (melting
point 262 K, boiling point 374 K; Weast, 1969) may
be responsible.
The surface temperature of Venus (c.74o K) is
close to the range of temperatures at lO7 Pa for agroup of decarbonation reactions. Three reactions
involving dolomite and various silicates are shownin fig. 9 (curves 4, 5, and 6), and many otherreactions, together with a thorough description of
the complexities of the phase relations are given in
Winkler (i 974). Reaction of spurrite + rankinite tolarnite + CO2 (curve 3) marks the upper temperature stability of calcium carbonate silicates.
The meionite variety of scapolite has a wide
stability range at high temperature and pressure in23

dry bulk compositions rich in the anorthite component (Goldsmith, r976; Newton and Goldsmith,I975). Its stability field is bounded by reaction toeither anorthite+calcite, or grossular+kyanite+
quartz + calcite, or liquid + crystalline phases +CO2. Meionite is only one of the components ofnatural scapolites on Earth, but exemplifies thehigh pressure and temperature to which thismineral group is stable. Scapolite occurs in deepseated granulite rocks and could trap CO2 escaping
from the Earth's mantle. Nitrate-scapolite (Goldsmith, Newton, and Moore, 1974) and nitratecancrinite are also stable at high P and T, but
probably do not occur on Earth and on otherplanets because the oxidation state is too low.
However, sulphate-scapolite (Goldsmith, i976 ) isstable on Earth (e.g. in the Sittampundi Complex
and in Australian granulites, Lovering and White,t 964), and must be considered as a possible storagesite for sulphur in other planets.
The eruption of carbonatite and kimberlite rocks
on Earth testifies to the storage of CO2 in the upper
mantle, and comparison of figs. 2 and 9 shows thatsimilar storage is theoretically possible on otherplanets if their bulk compositions are suitable.
Under reduced conditions (Brey and Green, 1977;
Rosenhauer et al., 1977) , carbon can exist as either
graphite or diamond, and occurrence of thesepolymorphs in mantle-derived xenoliths is consistent with the experimental transition curve(Boyd, 1973; Dawson and Smith, I975; McCallumand Eggler, 1976 ). In the deep mantle carbon might
be envisaged as substituting for silicon in tetrahedral coordination, but this has not been demonstrated in natural minerals. Cohenite, (Fe,Ni)3C,
occurs in iron meteorites, but the conditions for its
stability are uncertain (Anders, I964) , and do not
require high pressure as was once believed.
The occurrence of sulphur in the planets isgoverned by many complex phase equilibria. Progressive condensation of the solar nebula should
give large amounts of FeS (troilite), which mineralis indeed present in iron meteorites and as anaccessory in other meteorites and lunar rocks.
Sulphur is an obvious possible constituent ofmetal-rich cores for the terrestrial planets, as
proposed for Earth by Murthy and Hall (197o).
Small amounts of Fe,Cu,Ni-sulphides occur inmantle-derived xenoliths (Bishop et al., I975), andsulphur dissolves in small amounts (o.n wt ~o) in
mafic and ultramafic melts (Haughton et al., 1974;
Shima and Naldrett, i975). From electron microprobe analyses, Anderson (1974) found o.oo5-o,28
Wt~o S in glasses trapped inside megacrysts of
volcanic rocks from Hawaiian and circum-Pacificareas, whereas vesicular glasses contained only
o.oo5-o.o4 wt~o. Apparently the magmas were

24

J . V . SMITH

saturated before eruption with respect to a sulphidemelt, and lost considerable SO2 during eruption.
The geochemical distribution for sulphur is very
complex, involving emission of volcanic gases into
the atmosphere, and transport into aqueous andsolid reservoirs, and probably some recycling backinto volcanic environments. A review of the phaserelations in the system (Na, K,Mg, Ca).(C1,SO4) by
Rowe, Morey, and Zen (1972) points up the manypotential ways in which sulphate can exist near coolplanetary surfaces (e.g. of Mars). Numerous papers
cover the complexities of S-bearing ore-forming
fluids of which Holland (1965) and Naldrett (1969)
are representative. Anhydrite is stable to highpressures and temperatures in systems of theappropriate composition (e.g. occurrence in S.W.
African gneisses, Nash, I97Z), and sulphate-meionite is stable to even higher pressures and temperatures when the anorthite component is available
and the system is dry (Goldsmith, 1976, Fig. 9).
Formation of sulphate component in scapolite wasascribed to oxidation of sulphide by concomitant
reduction of F e 2 0 3 in iron oxide to FeO component in silicate. Sodalite and cancrinite minerals
are stable to high temperatures: thus sulphatenosean dissociates at 1528 K (Tomisaka andEugster, I968).
Chlorine and fluorine are further elements whose
occurrence at planetary surfaces may be indicativeof the internal composition. Anderson (1974)
observed o.oi-o.22 wt ~o C1 in the aforementioned
glassy inclusions, and concluded that as water boils
away from an erupting lava it carries C1 with it.
Hards (I976) showed experimentally that C1 tends
to concentrate in an aqueous vapour during differentiation of acid and alkaline magmas whereas F
tends to remain in the silicate melt and reduce thesolidus and liquidus temperatures. Wellman (i 97o)
described phase relations involving C1 in synthetic
syenites, and Tomisaka and Eugster (t 968) showedthat chloride-sodalite dissociates at 15o3 K. Concentrated brines can deposit halite and sylvine aspart of the group of evaporite minerals, and whenisolated these minerals melt at lO73 and IO43 K,
respectively. However, evaporite minerals are very
mobile and reactive in metamorphic terrains, andwould react with silicates to form feldspathoids atlower temperatures. Small amounts of C1 occur inamphibole and mica. Fluorine is strongly stabilized
in apatite, mica, and amphibole minerals (Ekstr6m,
I972; Paul, Buckley, and Nixon, 1976; Prins, 1973;
Stormer and Carmichael, 1971 ).
Space forbids similar surveys for other elements,
but a few comments are perhaps necessary forlarge-ion lithophile elements, which concentrate in
the late liquids differentiated by melting of basaltic
and peridotitic bulk compositions. Potassium isfound in diopside from mantle-derived peridotites(probably from c. 15o 2oo km depth) up to o.2 wt ~o(Erlank, I973; Sobolev, x972; Bishop et al., 1978)
and in pyroxene inclusions from diamond up too.9 wt ~, (Prinz et al., I975) but not so far in garnet(Bishop et al., 1978). Whether substantial potassium is stored in the deep mantle, and if so how, is
not known, but perovskite should be tested as apossible host. Phlogopite is a major host on Earth(Carswell, ~975; Smith and Dawson, I975), probably as deep as cA5o km and perhaps even to
30o km. In and near the crust, feldspar, mica andamphibole (Best, 1974) are major hosts, while some
evaporite minerals store potassium close to thesurface. Occurrence of K in metal cores of planets
would provide an attractive heat source for convection, but the chemical relations are controversial(summarized in Ringwood, 1975, p. 553). Sodium
occurs in pyroxene from mantle-derived xenoliths
up to 8 wt ~o and in garnet up to o.2 wt ~o (Sobolev
and Lavrent'ev, 1971; Bishop et al., I978 ). In thecrust feldspar is the dominant host but many otherminerals carry some sodium. If N a and K are
depleted in the Earth to the same degree, much of
the Earth's Na must occur below 2oo km depth(Smith, I977). In meteorites Na-rich plagioclase is
the dominant alkali species but several otherminerals occur of which the large K-feldspar
crystals in the Colomera iron and the Na-bearing
phosphate minerals are particularly interesting.
Isotopic data reviewed by Moorbath (1975) andSun and Hanson (I975) show that Sr and Pb havebeen stored in the Earth's mantle all throughgeologic time with intermittent eruption of primaryalkali basalt and nephelinite. These and other data
show that on Earth the large-ion-lithophile elements did not move entirely into the surface andthat mineral reservoirs exist perhaps down to atleast 6oo kin. Whether the reservoirs are independent minerals, or whether they are merely
microscopic inclusions trapped in stable minerals
produced during primary differentiation is notknown.
The simplest assumption for the terrestrial
planets is that the temperature did not reachvapourization temperature and that the only loss ofmaterial was escape of light species from theatmosphere.
Ringwood (1966a), in a review of the genesis ofchondritic meteorites, interpreted ordinary andenstatite chondrites as the possible products ofvolatile loss and autoreduction at high temperatureof primitive material like that of C I meteorites. He
emphasized the effects of selective volatility and of
selective solubility in supercritical fluids of H20,
CO2, H2S, and H2. Numerous studies of lunar
basalts and breccias demonstrated the mobility ofM I N E R A L O G Y OF THE PLANETS
C and S; thus Gibson and Moore (i973) found that
I2 ~o 3~ ~o of the C and S in lunar breccias could bevolatilized under vacuum in I day at IO23 K andalmost all at I373 K. Smith and Steele (I976)
reviewed the pertinent lunar mineralogy withrespect to chemical processes involving C, P, and S.
Particularly relevant to the chemistry of the earlysolar system are the studies of implantation of thesolar wind into fine-grained regolith and the subsequent effect on metamorphism. In general theseprocesses lead to reduction with removal of Fe
from silicate, phosphate, and sulphide, and formation of iron metal, schreibersite, and probably
cohenite.
Carrying further his preference for carbonreduction in a condensed environment (e.g. aplanetesimal) rather than hydrogen-reduction in a
low-pressure nebula, Ringwood in a series of
papers (see Ringwood, r975) developed the singlestage chemical model for accretion and differentiation of the terrestrial planets, particularly the Earthand its Moon. High temperatures result from rapid
accretion in which accretion energy overwhelms
radiation loss at the surface, and from formation of
an iron-rich core, which releases gravitationalenergy. The inner region of the solar nebula would
condense to oxidized assemblages of minerals likethose in CI carbonaceous meteorites (including
magnetite, hydroxylated silicates, and carbon-richcompounds) while the outer region would bedominated ultimately by ices of various compositions including methane and ammonia. Ringwood
developed the terrestrial planets by rapid accretion( < Io 6 years) before the T Tauri wind began. Such
rapid accretion produced temperatures exceeding
13oo K during most of the growth of the Earth, andfrom I8oo to over 2300 K at the end. A cool,oxidized, volatile-rich nucleus of olivine, magnetite,hydrated Fe,Mg,Ni-silicates, and (Fe,Ni)S, withsurface temperature near Iooo K, became reducedby carbonaceous material and degassed as further
accretion to about one-fifth of final mass brought
the surface temperature up to 15oo K. A primitive
reducing atmosphere developed from the reaction
CH4 + H 2 0 ~ CO + 3H2 . Further accretion withrise of surface temperature to I8oo K caused
reduction and volatilization of some relatively
volatile elements including alkali metals, Zn, Hg,
Pb, S, and C1. Accreting material reacted with theatmosphere to give metallic iron and Fe-free silicates. Further temperature rise to over I8oo K
caused volatilization of SiO and accretion of
forsterite. At even higher temperature Mg vapour
would be formed and metallic Si would enter iron
metal. Calcium and aluminium silicates would
accrete. Gravitational instability with sinking of
metal to form a core would yield 2.5 J/kg, and the25

entire outer part of the Earth would volatilize into a
disc. Condensation into a sediment ring that
accreted into a Moon, and removal of the volatiles
by an intense T Tauri wind complete the singlestage model. This impressive model has many
attractive features but has been criticized. How canenormous amounts of CO be lost from the Earth?
How can substantial amounts of SiO be volatilized
from the Earth? Some modifications to the singlestage model are described in Ringwood (I975, pp.
59I-7). I prefer to consider models that do not
require escape of substantial amounts of heavy
molecules from planets with escape velocities
greater than 5 km/s.
Chemical differentiation in a planetary atmosphereand consequent effect on the solid part. Escape of anatom, molecule, or ion can occur from a planetaryatmosphere when the escape velocity is exceeded.
Simple kinetic theory of gases leads to a classicaldistribution of velocities whose mean value for an
ideal gas is ~" -- (8 kT/~zm)~, where k is Boltzmann's
constant, T is temperature, and m is mass. Escapeoccurs from the high-energy tail of the distribution.
Unfortunately, the loss from planetary atmospheres is complicated by the variation of temperature and chemical composition with height, theextent of ionization from solar radiation (principally ultraviolet), and the effect of the solar wind. At
one extreme the T Tauri wind has been used to
sweep away the entire atmosphere, irrespective of
mass (Ringwood, t975), while at the other extreme,
the upper atmosphere might act as a cold trap
inhibiting or preventing the loss of all but thelightest species. Volatiles can be ejected to highaltitude by volcanic emission, and removed byadsorption in soil or by direct freezing. It is difficult
to place fundamental constraints on the escapeprocesses, and readers are referred to Hunten(I973) , Hunten and Donahue (i976), Curtis andHartle (I977), Visconti (I977), and articles listed in
Mars and Venus sections.
It is convenient here to mention Turekian's latest
model for the Earth, which depends greatly on lossof volatile species. A brief summary (Turekian,I976) mentions the following features: hot condensates accreted sequentially and rapidly on to thegrowing Earth, followed by a rubble veneer, whichbecame progressively enriched in low-temperature
products; the rubble veneer did not equilibrate with
the iron-nickel core; most of the iron from thenebula was sequestered in large bodies duringcondensation and did not react with the remaining
gas; low-temperature condensates in the rubble
veneer lost water to the atmosphere during the firstI5 x IO6 yr, and most of the protons in the water
were lost from Earth, leaving oxygen which combined with metallic iron to give the Fe 2 + and Fe 3 +26

J. V. SMITH

in the present crust and upper mantle; after dissipation of the early atmosphere the atmosphere would
be primarily composed of H 2 0 and CO2 with theminimum oxygen pressure determined by theoxidation state of the crust and upper mantle;
oxidation of Fe metal to FeO resulted in conver-sion of some pyroxene (c.MgSiOa) to olivine(Mg, Fe)2SiO4.
Instead of following Turekian's model, thealternative of losing most of the H from planetesimals should be considered.

EARTH
Descend into that crater of Sneffels Yocul... audacious traveller, and you willreach the centre of the earth. I have done it.
Jules Verne, Journey to the Centre of the Earth
We can only envy Arne Suknussem in our puny
attempts to explore the interior of the Earth!
This section is restricted to mineralogical, petrological, and geochemical properties particularly
relevant to the general mineralogy of the planets.
Speculations on the earliest history of the Earthrange from complete melting and partial volatilization, even of Si, to slow accretion with segregationof a core and proto-crust by localized partial
melting and diapirism. The latter approachpermits many variants including chemically homogeneous vs. inhomogeneous accretion, and rapid
vs. slow differentiation into a core, mantle, andcrust. Emphasis is placed here on the relics of theoldest crust, and on the upper-mantle xenoliths
carried up by alkali basalts and kimberlites.
General reviews of the properties of and processes
inside the Earth were given by Stacey (I977),
Ringwood (i975) , and Wyllie (I97i).
Early history

The accretion age of the Earth is hard to define,and all estimates depend on geochemical models.
An estimate of 67 _+2o million years for the formation interval of the Earth (Pepin and Phinney, 1976)
is consistent with interpretation of initial Pbisotope data for the Amitsoq gneiss, W. Greenland(Gancarz and Wasserburg, I977).
Advances in the interpretation of structuraldeformation in polymetamorphic rocks are leading
to decipherment of early Archaean terrains, assummarized by contributions to a N A T O
Advanced Study Institute (Windley, I975).
Comparison with data for Proterozoic terrains(--2. 5 to - o . 6 x I O 9 yr) and later assemblages
allowed Windley (I 977) to summarize the evolutionof the continents. As far as possible references are
made to articles by listed authors in Windley (1975).
That the oldest known crystallization age for aterrestrial rock is about - 3 . 8 x Io 9 yr (Moorbath)is not surprising in view of the lunar evidence for
bombardment by major projectiles (2o-loo km

diameter?) as late as - 3 - 9 x l O 9 yr. Coupled withthe trebled production of radioactive heat in the
pre-Archaean era, the disruptive effect of the
projectiles should have produced major volcanic
activity in the outer part of the Earth. There is noneed to invoke substantial accretion of material
after - 3 . 8 • Io 9 yr, though minor amounts havearrived as meteorites, comets, and dust.
Petrologically the Archaean rocks can be separated into two metamorphic suites. A high-grade
one is dominated by tonalitic-granodioritic
gneisses of calc-alkaline affinity, intercalated withlayered igneous complexes and metamorphosed
volcanic and sedimentary rocks of which marbles,graphite-bearing mica schists, and banded magnetite-quartzites may be particularly significant. A
low-grade suite of greenstone belts was originallycomposed of volcanic rocks of ultramafic, mafic,
and calc-alkaline affinity, and sedimentary rocks
including cherts and banded iron formations.
Attempts to compare the chemistry of theserocks with modern ones are complicated by theprobability of extensive metasomatism (Gunn), assuggested by the depletion of mobile elements (e.g.
K, Rb) in granulites compared to amphibolites, andby the certainty that higher heat flow would resultin a different tectonic style (Burke, Lambert) and indifferent processes of partial melting in Archaean
than in modern eras. Both Archaean suites containrelics of ultramafic volcanic rocks, including
komatiites characterized by spinifex-textured
Mg, Cr-rich olivine. Such olivine-rich rocks are
relatively less abundant in later eras, which supports ideas that the higher heat production in the
Archaean era allowed a greater degree of fractional
melting (Green, I975).
Since blue-schists are virtually absent in Precambrian rocks, while green-schists are abundant backto the Archaean, plate tectonic processes must havechanged considerably with time, if indeed suchprocesses can be applied to the Archaean era.
Bearing this in mind, Windley and Smith noted theresemblance between the rocks in the Archaean

MINERALOGY OF THE PLANETS
high-grade terrains and those in active continental
margins along Cordilleran fold belts, while Tarney
et al. used the 'Rocas Verdes' marginal basin as amodel for greenstone belts. Windley (I977, pp. 6o3) incorporated these features into a plate-tectonic
model for the Archaean suites. Even more speculative are the models for a pre-Archaean 'scum'dominated by basic rocks (Lambert, Murthy,
Shaw), which turned into proto-continents via a'permobile' tectonic regime (Burke). Perhaps Carich amphibole and Na-rich plagioclase dominatedthe proto-crust in contrast to Ca-poor pyroxene
and Na-poor plagioclase in the lunar crust.
Interpretation of Rb-Sr and U - P b isotopic data
depends considerably on models in which the rateof, and extent of, separation into separate reservoirs by partial melting and metasomatism provide
difficult problems. Preliminary investigations(Moorbath) imply that Archaean and Proterozoic
gneiss complexes, greenstone belts, and 'granitic'
plutons are mostly new additions to the continental
crust by episodic partial melting of the mantle,
while regeneration of the crust, except for rocks
rich in K-feldspar, is a minor process. This implies
that feldspar-rich rocks are irreversibly differentiated from the mantle with little returned by
subduction. Models for the origin of the Earth
Moon system must allow for storage of reallysubstantial amounts of large-ion-lithophileelements in the Earth's mantle for at least IO9 yr
after accretion. These ideas are compatible with theconclusion by A. T. Anderson (i974) that there is a
gradual and permanent transfer of C1, S, and H20from the interior of the Earth to surface reservoirs
on a time scale of the Earth's life.
Clues to the development of the atmosphere andhydrosphere have been inferred from reconnaissance studies of the sediments. From the relation
between sedimentary and volcanic rocks in Archaean terranes, Sutton suggested deposition of
sediments on stable shelves. The stratigraphic
succession for the Barberton greenstone belt has
ultramafic-mafic volcanic rocks overlain by calcalkaline volcanic rocks, and finally by sediments
ranging from a lower argillaceous group of
shales, sandstones, and greywackes to an upper
arenaceous group of conglomerates, quartzites,
limestones, and banded iron formations. These
sediments can be assigned to a range of deepto shallow-water environments (see also Lowe andKnauth, I977).
The extreme depletion of noble gases on Earthrules out direct capture of gases from the solar
nebula, and all recent models assume that thepresent atmosphere was derived by out-gassing of
the interior, modified by an unknown loss of lightgases through the atmosphere and ionosphere.

27

Present volcanic gases are composed of roughly
equal amounts of H20, CO2, and SO2, possibly inoriginal equilibrium with magnetite-bearing rocks,
and certainly diluted during ascent by erratic
amounts of contaminants (Nordlie, I977).
Current theories (Schidlowski, Walker) assume
that early volcanic gases were dominated by H 2 0
and CO2, and indeed the sedimentary and volcanic
rocks dating back to - 3 . 8 x io 9 yr have beeninterpreted in terms of an early HzO-COz-richatmosphere, essentially free of 02. Loss of H
through the ionosphere was coupled with captureof the released Oz by Fe z +, C, and S. Free oxygen
developed in the atmosphere by fermentation produced by organisms, which became abundant in theProterozoic era. However, it is not clear what was
the temporal variation of oxygen in the atmospheresince - 3.8 x IO9 yr, and what was the nature of theatmosphere before - 3.8 • IO9 yr, especially during
the accretion of the Earth. Whether methane andammonia were present in the early atmosphere is of
considerable biological significance because of the
speculations on the origin of amino acids and otherorganic molecules.
Some quartzites contain abundant magnetite inbands, and these banded iron formations, datingback to - 3.8 x l O 9 yr, typically occur with volcanic
rocks, perhaps consistent with derivation of theiron and silica during fumarolitic activity, as suggested by Goodwin (Windley, 1977, P. 42). Alternatively, the iron and silica may have derived fromweathering of continental rocks in a reducing
atmosphere such that iron was washed away as
Fe z+ rather than staying as Fe 3+ as in modern
sediments. Schidlowski discussed Cloud's suggestion that autotrophic organisms were protectedfrom ultraviolet radiation by an aqueous shield,
and released O2, which combined with hydrated
Fe z + ions to ultimately yield magnetite in sediments. Arguing against the weathering model is theabsence of A1203-rich residues in Archaean continental relics, which would be expected after leaching of Fe z+ (Windley, 1977). Perhaps a combination of fumarolitic activity with biological fixation
provides a viable model. Whatever the source of thebanded iron formations they could act as a sink foroxygen produced by the exospheric loss of H fromatmospheric water. Banded iron formations became sporadic after - 2 • I O 9 yr, and continental'red beds', closely associated with evaporites, became abundant thereafter. Schidlowski suggested
that these hematite-bearing rocks began to form
when oxygen was released to the atmosphereresulting in conversion of weathered Fe 2+ into
immobile Fe 3+. Perhaps a slow-down of fumarolitic activity might be involved. Ultimately theconcentration of oxygen in the atmosphere reached

28

J.V.

the 'Pasteur' level of one-hundredth the present
level, at which stage breathing organisms
developed. Fryer (I977) observed anomalously
large concentrations of Eu in chemical sediments
earlier than - 1.9 x IO 9 yr and anomalously high Ce
for those later than - 2. 3 x 109 yr. This was interpreted as the result of increasing oxidation suchthat Eu 2+ and Ce 3 + were stable during formationof the earlier sediments and Eu 3+ and Ce 4+ for thelater sediments.
Graphite is widespread in metasedimentary
Archaean rocks including mica schists, limestones,
and banded iron formations such as the - 3.8 x IO 9
yr Isua occurrence (Allaart). Carbon isotopic ratios
indicate that a sedimentary reservoir of organiccarbon had built up by - 3.7 x i0 9 yr, indicating
early existence of photosynthesis (Schidlowski).
Morphological evidence of early biological activity
is controversial (cf. Schopf vs. Muir and Grant).
The ratio of C/CO2 in sedimentary rocks has beenroughly constant over the past 3 x lO9 yr. Holland
suggested that decreasing heat production slowed
down geochemical cycles, and in particular lowered
the injection of CO2 and reducing gases into theatmosphere. Carbonate rocks date back t o - 3 . 8
x io 9 yr, with a temporal trend from dolomite-richvarieties in the Archaean to calcite-rich ones inmodern sediments (Holland).
Opposing the concept of an early oxygen-poor
atmosphere, Kimberley and Dimroth interpreted
uraninite and pyrite in Archaean sediments as
products ofdiagenesis rather than as indicators of alow partial pressure of 02 in the atmosphere ( < 4x io -3 bar: Holland). They ruled out radiativesynthesis of hydrocarbons in the atmosphere because of the absence of bitumen in sediments, andnoted that carbonaceous mudrocks of all ages arerich in pyrite. Finally they explained the absence ofArchaean red beds and sulphate deposits by lack ofpreservation of appropriate shelf facies. Simpson
and Bowles (i977) also suggested that the earlyatmosphere was non-reducing because of evidence
for SO 2- and UO22§ ions in solution; furthermore
they argued that uraninite grains of the Witwatersrand and Dominion Reef Systems were detrital,and had survived abrasion because of hardening
from Th substitution.
Obviously there is great scope for study of all
Archaean rocks, perhaps especially of sediments.
For comparative planetology it seems safe toconclude that by - 3.7 x l O 9 yr the Earth had a nonreducing atmosphere, perhaps even containingsome molecular oxygen, and that proto-continents
and shelves were becoming sufficiently stable tosurvive later tectonism (cf. Hargraves, ~976).
Furthermore, the upper mantle was spewing out
volcanic rocks mostly similar in chemistry to those

SMITH

of recent times, except perhaps for absence ofkimberlitic and carbonatitic types thought to be
characteristic of large stable cratons, and forpresence of the ultrabasic types indicative of agreater degree of partial melting (e.g. Villaume andRose, I977). In detail, there is evidence of mantle
heterogeneity and chemical evolution with time(Sun and Nesbitt, I977), but for this general review
it is probably better to point to resemblances rather
than differences between Archaean and modern
crustal rocks. Speculations that the Earth had amassive primitive atmosphere (e.g. 3o0 bars H20
and 50 bars CO2; Walker) are not ruled out by
direct evidence from Archaean rocks. However,there is no evidence in support of proposals for a
massive reducing atmosphere, dominated bymethane, from which organic molecules could be
generated. It is necessary to turn to evidence in theupper mantle in order to proceed further with amineralogical study of the Earth.

Upper mantle (above 200 km depth)
Kimberlite magmas have formed in stable continental regions at depths up to c:15o-2oo km, andduring explosive emplacement have torn off andtransported fragments of the wall rocks--hence thename 'Poor Man's Mohole' for a kimberlite pipe.
Basalt magmas provide information on the site ofpartial melting, and alkali basalts have transported
pieces of the crust and mantle down to c.6o km. The
details are very complex (e.g. Ringwood, 1975;
Meyer, 1977; Sobolev, I977), but an over-all picture
has emerged on the basis of the following eight
concepts: the dominant rock types of the upper
mantle are Ca, Al-rich peridotites, Ca, Al-poor harzburgites, and subordinate eclogites and dunites.
The peridotites are dominated by olivine andorthopyroxene and are subdivided into spinel- andgarnet-varieties, which can be attributed respectively to approximate depth ranges of c.25-c.6o km,
and > c.6o km when the oceanic and continentalshield geotherms are compared with a generalized
phase diagram for peridotites (fig. 2). Partial melting of Ca,Al-rich peridotites would result in extraction of various basaltic magmas as the Ca,Al-richminerals were dissolved, leaving behind refractory
residues dominated by olivine and orthopyroxene,
thereby explaining most harzburgites. Althoughconsiderable amounts of Ca,Al-rich material haveentered the crust as basaltic regions near the surface
and metamorphosed amphibolitic to granulitic
regions at depth, large amounts remain in the upper
mantle. There is a very large range of compositions
for the pyroxenes and garnets in eclogites andgarnet peridotites, and diverse origins are needed
including metamorphism of arrested partial melts

M I N E R A L O G Y OF THE PLANETS
and of subducted rocks. There may be a tendency
for density differentiation with heavier fertile
peridotites underlying lighter barren peridotites(Boyd and McCalli~ter, 1976). [Note: 'fertile' refers
to the chemical capacity to yield basalt upon partial
melting, as proposed by M. J. O'Hara.] There isstrong evidence for substantial heterogeneity of thesource regions of basalts, though the details are
based on controversial interpretations of geochemical data (e.g. Hofmann and Hart, i978; Kayand Hubbard, 1978; Tatsumoto, 1978 ). Finally,
there is no need to assume major changes in thechemistry of the source regions of basalts sinceArchaean times (O'Nions and Pankhurst, 1978 ).
The details of upper-mantle mineralogy are
extremely complex, as is apparent from the Proceedings of the First and Second International
Conferences on Kimberlites (Phys. Chem. Earth, 9;
to be published), and many papers scattered among
journals. Although most olivines and orthopyroxenes from upper-mantle rocks have
Mg/Fe e.IO, the contents of trace and minor elements are highly variable; e.g. harzburgites can bedivided into 'barren' and 'fertile' varieties on thebasis of Ca and A1 substitutions in the orthopyroxene (Hervig et al., 1977). Major and minor
elements in garnets and pyroxenes vary considerably in the upper mantle, and cluster analysis isyielding classification schemes (Dawson andStephens, i975; Stephens and Dawson, 1977) thatshould lead to petrogenetic interpretations.
Controversy remains over the experimental andtheoretical calibration and the practical application of the pyroxene and garnet barometer-thermometers for garnet peridotites (many papers), but thegeneral concepts and the application to estimationof a shield geotherm (Boyd, 1973) mark a high point
in upper-mantle petrology. In future years trace
and minor elements should become increasingly
used in thermo-barometry, and two possible candidates are the Na distribution between garnet andpyroxene (Sobolev, 1977; Bishop et al., I978), andthe Cr content of silicate and oxide minerals(Malinowsky and Doroshev, I977).
The roles of volatile elements are particularly
intriguing. Depression of the melting temperatureof peridotites by HzO and COz has been exploredin detail (see earlier). Under terrestrial shields, it appears that the maximum depth for stability of amphibole is c.Ioo km and for phlogopite is c.2oo km,
though this latter value might possibly be too
low for an unusually cool region. These minerals
should be major reservoirs for H 2 0 (actually asOH) in the upper mantle. Whether substantial H 2 0
can be stored at greater depths is totally speculative, but synthesis of hydroxylated Mg-silicates(Akimoto et al., 1977) is indicative. O f course, it is19easy to postulate hypothetical hosts, such as inclusions trapped in refractory minerals during accretion, and indeed the emission of primordial noble
gases from the mantle testifies to deep-seatedreservoirs that have persisted for 4.5 x io 9 yr (Craig
and Lupton, i976; Dymond and Hogan, 1978 ).
However, it seems easiest to assume that an early
differentiation flushed out most of the water fromthe mantle below c.2oo km. A similar conclusion isprobably true for CO2, but again there is experimental evidence indicative of potential incorporation of CO2 into crystalline material at lowermantle pressures (Jeanloz and Ahrens, I977b ).
Absolutely crucial to the controversies over theaccretion of the Earth are the contents of chalcophile and siderophile elements in the upper mantle(Smith, I977) , which indicate core-mantle disequilibrium, as emphasized by Ringwood (I966b). The
Fe/Ni ratio of upper-mantle olivines is close to thecosmic ratio, thereby indicating that Ni was not
stripped away significantly from the upper mantle
by reaction with either metallic iron or sulphide at
low pressure. Silicate sulphide partitioning (Binns
and Grove, I976; MacLean and Shimazaki, 1976;
Rajamani and Naldrett, 1978 ) depends on complex
factors, but the partition coefficients for Ni favour
sulphide so strongly that it must be concluded thatsulphide is only a minor component of the upper
mantle, even though it is a ubiquitous trace constituent of peridotite and basalt. The noble metals in
peridotite xenoliths (Smith, 1977) are strongly
depleted over cosmic composition, but the depletion factors are not as high as would be expected forequilibration of peridotite with metal or sulphide(Kimura et al., I974; Rajamani et al., 1977). These
data are strong evidence for late accretion ofmaterial either undepleted or only slightly depleted
in chalcophile and siderophile elements, and it may
be possible to go even further and argue that the
Earth then had a stable mantle barrier to prevent
equilibration with a metal-rich core.
Diamonds have a triple significance: pressure
indication (fig. 9), redox indication (Brey andGreen, 1977; Rosenhauer et al., I977), andpreservation of inclusions. Existing data indicate
growth in both peridotitic and eclogitic hosts, andthe possibility of survival of diamond during
partial-melting events should be considered. Of
great potential importance are the observations ofhigh values of K 2 0 in some pyroxene inclusions(Sobolev, 1977) and of the very high content of Cr
in spinel inclusions (Meyer, I977). Although diamonds have been found in situ in eclogite andperidotite xenoliths, the silicate and spinel inclusions of random diamond grains mostly do notmatch in detail the chemistry of silicate and spinelminerals of the xenoliths, and it may turn out that30

J . V . SMITH

the inclusions 'remember' an earlier igneous event
whereas the independent minerals are mostly controlled by metamorphism, perhaps following an
earlier igneous event.
The trace-element composition of the upper
mantle is usually inferred from partial-melt
models applied to observed chemistry of basalts.
Since partial-melt models involve complex
assumptions about the melting process and theextent of crystal-liquid differentiation duringmagma ascent, they are not pursued here. For
general reviews readers are referred to Wyllie(~971), Carmichael et al. (1974), and Yoder (I976),
and for more recent papers to references listed by
Sun and Nesbitt (i977) and those in the Memorial
Volume to P. W. Gast (Earth Planet. Sci. Lett. 38,
no. I). Ideally, this approach should be tested byanalysis of peridotite and eclogite xenoliths. Unfortunately, metasomatism during transport to thesurface causes systematic chemical bias of bulk
samples, and it is desirable to proceed to chemical
analysis of unaltered regions. Most existing bulk
analyses of mineral separates are affected bymechanical impurities, as demonstrated by highconcentrations of incompatible elements in olivine,
but the data of Shimizu (i975) for K, Rb, Cs, Sr, andBa in clinopyroxenes from ultramafic xenoliths
appear especially good, probably because of carefulhand-picking of clean grains. Thorough attention
to the background is allowing detection levels of
c.2o ppm for many elements measured with anelectron microprobe, and Smith et al. (1978) wereable to compare K/Rb/Ba ratios for micas ofseveral textural types in peridotite xenoliths andkimberlite host with ratios for various types ofbasalts. Micas with primary texture in peridotites
have K/Rb close to that for CI meteorite (c.3oo),
while K/Ba is about six times lower. Many basalts
could have obtained their K, Rb, and Ba contents
by total melting of mica from peridotite, whereas
those with high K/Rb (c.Io 3) could not. The next
decade should see a plethora of data on traceelements in mantle-derived minerals using ion
microprobes, as indicated by measurements oftransition elements (Shimizu and All6gre, I977),
thereby placing limits on models for partial melting
and metamorphism when matched with experimental data for element partitioning (see especiallythe recent Year Books of the Carnegie Institution of
Washington).
Because the upper mantle is heterogeneous as aresult of partial melting and mechanical transport,
it is not surprising that estimates of the bulk
composition vary considerably. Analyses of garnet
peridotites range several-fold in Ca and A1, probably as a result of partial melting, and attempts
have been made to select compositions supposedly

undepleted or unaugmented in basaltic component.
Other attempts involve addition of enough basaltic
component to peridotite to give CI-meteorite
ratios of rare earths. Some analyses of garnet
peridotites are listed in Smith (1977, Table 4).
Ringwood (1975) obtained a model 'pyrolite' byadding one part of 'basalt' to three parts of'peridotite', and this is a useful but not definitive
composition for speculation about the mineralogyof the Earth's mantle.
Mantle (deeper than 2oo kin) and core

All mineralogical speculation for these regions
depends on interpretation of seismic evidence interms of a density-depth profile (fig. I) and velocity-depth profiles (not shown), cosmochemically
based speculation on the bulk composition, andexperimental measurement of the physical properties of likely minerals. A major uncertainty is thetemperature profile (McKenzie and Weiss, 1975),
which is constrained only weakly by the liquidsolid transition in the core. Thomsen 0977) questioned the theoretical foundations of equations ofstate for terrestrial planets, and the present review
will emphasize the uncertainty of present interpretations.
The properties of the core (reviewed by Brett,r976, and Jacobs, i975) are important for speculation on the bulk composition of the Earth (nextsection). Although the solid inner core must be richin iron and siderophile elements (particularly Ni),the chemical details are unknown. The liquid outer
core is less dense than an Fe,Ni liquid (fig. 1), but thelight substituents cannot be uniquely defined.
Troilite or pyrrhotine occur in most meteorites, andpartial melting of chondritic compositions would
give heavy Fe,FeS-rich liquids which should sink ina growing planet. Thus the suggestion that S is the
major light element in the Earth's core (Murthy andHall, I972; Murthy, 1976) is very popular, andAhrens (1978) concluded that 9-I2 Wt~o S issufficient to explain the lower density. Ringwood
originally argued for Si, by analogy with Si-bearing
metal in highly reduced enstatite meteorites, but
has now (1977) switched to oxygen because of twoarguments: first, placing c. I O ~o S in the core would
lead to greater depletion of K (six-fold) than S(three-fold) with respect to solar composition,
which is contrary to the Grossman-Larimer type ofmodel for simple progressive condensation of thesolar nebula, and second, thermochemical calculations indicate complete miscibility between Fe
and FeO above 38oo K at high pressure. Smith
0977) argued that S might be protected frommechanical disaggregation in Fe-metal cores ofdifferentiated planetesimals while K was lost

M I N E R A L O G Y OF THE PLANETS
preferentially from the bombarded surface. Currently, it is desirable to consider the two extremes of
S-rich and O-rich outer cores. Simple analogy withiron meteorites suggests that C, N, and P may also
substitute in the Earth's core. Smith (I978) tentatively suggested that U and Th may partition into
troilite under reducing conditions, thereby providing a possible heat source for the Earth's dynamo,
but the available evidence from troilite nodules inmeteorites is poor, and probably unfavourable.
Whether substantial K enters a S-rich core is highly
controversial (Goettel, I974; Ringwood, 1975) andthe latest experimental partition coefficients between silicate and sulphide (Ganguly and Kennedy,1977) apparently place an upper limit of IOOppm K
in the Earth's core.
The mineralogy of the lower and middle mantle
is quite speculative. Several transitions are indicated by seismic studies, as exemplified by adensity-depth profile (fig. I). Refinement Of seismic
data may lead to minor revision of the profiles,
especially if subducted slabs actually penetrate to
great depth (c.7oo km). To a first approximation the
mantle can be regarded as an assemblage of closepacked oxygen ions with interstices occupied bycations, and the physical properties can bemodelled by equations which depend principallyon the mean atomic number Z and not on thedetailed crystal structure. This approximation isinadequate, and must be replaced by equations thatutilize data measured for specific crystal structures.
The early structural data obtained by Ringwood
and collaborators (collected in Ringwood, 1975) arebeing extended and essentially confirmed by new Xray data obtained in diamond-anvil cells. All thesilicates found in peridotite (viz. olivine, pyroxenes,
and garnet) transform into denser structures as thedepth increases. Particularly important is theperovskite structure, which is probably the major
carrier of (Mg, Fe)SiO3 in the lower mantle, andquite possibly of other components including Caand Al-bearing chemical units (Liu, I974, 1975,
i976; Liu and Ringwood, 1975; Liebermann et al.,
1977; Sawamoto, 1977; Ito and Matsui, ~978; Yagi
et al., I977, I978 ). Other important structure types
are garnet and fl-Mg2SiO 4 (upper mantle), periclase and ilmenite (middle mantle), and perhaps astructure with 8-coordinated Si (similar to fluorite?) in the very deep mantle (Jeanloz and Ahrens,
1977a).
There are four major difficulties in disentangling
the various chemical factors. First, the Mg/Si ratio
may vary with depth. Ringwood favoured volatilization of SiO from a very hot earth resulting ina mantle dominated by Mg2SiO4, whereas D. L.
Anderson favoured a mantle dominated by
MgSiO3. Jeanloz and Ahrens (I977a) suggested

31that olivine and pyroxene compositions have similar physical properties at high pressure, whereas
Sawamoto (~977) interpreted the lower mantle interms of pyroxene composition and perovskite
structure. Second, there is uncertainty in the depth
variation of Fe/Mg, ranging from a model with
substantial increase of FeO with depth (e.g. reaching 2I wt ~o in the deep mantle; Anderson, 1977) tothe possibility of a smaller change of FeO (e.g.
Davies, I974), and to the interpretation by Sawamoto (1977) that the ratio actually increases withdepth from o.o5-o.18 in olivine composition atzoo 35o km depth to o.I I-O.Z5 in olivine composition at 4oo-6oo km and then decreases to o-o.I8
in pyroxene composition at 7oo-I5OO km depth.
Third, there is essentially complete uncertainty ofthe spatial distribution of Ca, A1, Ti, and all otherelements of lower concentration, and fourth, disproportionation of 3FeO into Fe and Fe203 may
occur above 1oo kb (Mao and Bell, I977a ).
Undoubtedly these uncertainties will be reduced bydetailed crystallographic studies with diamondanvil devices (now operating up to 1.5 Mb, Mao
and Bell, I977b), phase-equilibrium studies to
2oo kb or higher (e.g. Akimoto and Akaogi, I977),
and crystalchemical models (e.g. Hazen andPrewitt, 1977; Burns, 1974) based on detailed crystal structure analysis (e.g. spinel at high pressure,
Finger et al., i977). To illustrate technical improvement, note how revision of the cell dimensions ofquenched MgSiOa-perovskite changed the densityfrom 4.oo g/cm a (Sawamoto, I977) to 4.1o (Itoand Matsui, i978; Yagi et al., r978), thereby allowing a lower Fe content for any part of the
mantle containing perovskite.
In the meantime Anderson (1977) suggested
transformation of pyroxene to garnet at 25o37o km depth, olivine to fl-phase(?) at c.4oo km,
and pyroxene to ilmenite(?) near 5oo km, whereas
Liu (I977) suggested that the 4oo km discontinuity marks the transformation of olivine to flphase and gradual absorption ofpyroxene into garnet, the 65o km discontinuity involves formationof perovskite, ilmenite, and periclase structures,and the lower mantle has perovskite replacing
ilmenite.
Bulk composition

Speculation on the bulk chemical composition,
especially its radial distribution, is important for
constraining models for the chemical composition
and mineralogy of the inner planets, and especiallyof the Earth Moon system. As emphasized inpreceding sections, it is impossible to model the
Earth merely from direct observations, and it isnecessary to use cosmochemical information.
Several attempts by geochemists to model the32

J . V . SMITH

Earth in the I96OS were followed by Larimer (I97i)
who compared twenty-two elements plus I - I 2 0 in
the Earth's crust with those in ordinary chondrites
and eucrites, by Shaw (t972, 1976 ) who developed adetailed model for chemical differentiation, and by
Clark et al. (I972) and Hutchison (i974, 1976),
whose models invoke two different kinds of heterogeneous accretion. This section is a heavily truncated version of Smith (1977), with some new ideas.
Rather than model the Earth with mixtures of
meteorites (Anderson et al., I971 ; Murthy and Hall,I972 ), Ganapathy and Anders (1974) assumed thatthe inner planets were made by 'exactly the sameprocesses as the ordinary chondrites'. Equilibrium
progressive cooling of the solar nebula at ~o 4 barwith C/O c.o.6 yielded the theoretical condensates
shown in fig. 4- Ganapathy and Anders split theelements up into five groups belonging to early
condensate, silicate, metal, volatiles condensing at
I3oo-6oo K [lithophile and siderophile subgroups], and volatiles condensing at < 6oo K [twosub-groups]. They assumed that each group was
unfractionated after condensation except for themelting event that formed chondrules. Relative
fractionations between groups were calculated byassuming Cl-meteorite composition for the element ratios in the nebula and the following Earthdata: o.oi8 ppm U to model the early condensate;
36 wt ~o Fe in the bulk Earth to give the metal/silicate ratio; observed Fe/Mn ratio to give oxidation
state of iron; assumed S content to give FeS; K/Uratio of 944 o to model loss ofvolatiles (I 3o0-6oo K)
during assumed chondrule formation; and T1/U
ratio of 0.27 to give volatiles ( < 6oo K). Brevity
forbids further details, and the final composition is
given in Table IV. Ignoring details the resulting
model has the Earth enriched in 'early condensate'and 'metal' about 1.5-fold with respect to Mgsilicate, and depleted about four-fold in 'volatiles
13oo 600 K' and about fifty-fold for 'volatiles< 6oo K'. Special treatment for H, C, and N
resulted in depletions of 5 x io-7(H), 5 x IO 4(C),and 3 x 1o-5(N).
Smith (1977) made further estimates of the bulk
Earth trying to utilize observed compositions as far
as possible. There was no problem for the atmosphere and hydrosphere. Important elements in thecrust were modelled after Larimer (i971), andothers were taken from compilations by Ronov andYaroshevskiy and by Taylor. Probably the major
uncertainty is in the lower crust, where high-grade
metamorphism may have caused substantial redistribution of mobile elements; however, there issubstantial agreement between a new estimate by
Ronov and Yaroshevskiy (1976) for the whole crustand the estimates by Leyreloup et al. (I977) for the
lower continental crust under a French volcanic

TABLE IV. Suggested chemical compositions of theplanets
wt

a

H
C
N

ppm
ppm
ppm

O

~o

Na
Mg
AI
Si
P
S

b

c

d

e

78
350
9.1

ppm

66
220
19
28"5 3 H
1 5 8 o 850

66
220
19
31.3
850

0.8
II
o.I
42.0
71o

---3 5 - 3 I5-5
300
--

~
~
~
ppm
~oo

I3-2I
1.77
I4.34
215o
1.84

14-65 13.7
1.3o 1.83
16.o9 I5.I
17oo 183o
4.34 2-9I

C1
K
Ca
Ti
Cr

ppm
ppm
%
ppm
ppm

25
45
I7o
13o
1.93 1.62
Io3o 66o
4780 3700

45
13o
2.28
928
416o

Mn
Fe
Ni
Zn
U

ppm
G
~
ppm
ppb

59~
35.87
2.04
93
18

470
~ 1 4 o lOO
31.7 II. 5 29.2
1 . 7 2 0.65 1.7
I47
3.i
20
40
(I2)

470
28.19
1.53
22o
14

f

g

-4 I-617.8 1 5 " ~ 5 - 3 4 20.5
3.92 1.2
3.I7 1.7o
19.2 1 6 . 4 6.26 19.2
69o (moo) 37oo
-57~
?
-o.47
76
3~
830
4.28 1.3
3-93 2.00
2300 700
I6OO
17oo (3800) 8000
61.1
3.4

13.o

34

--

a. Earth. Ganapathy and Anders (i974). b. Earth. S1 model(Smith, 1977) based on total condensation of metal, silicate andhigh-temperature fractions, and various depletions of volatileelements, c. Earth. New $2 model in which the S~ model is revised
to have Io ~ S in core, 'metal' elements augmented by 1.2, 'early
condensate' elements by 1.5 over (Mg + Si). d. Moon. Morgan e tal. (1978b); model 4b based on lowered U content and assumed
core mass of 2 % (mostly Fe, Ni, S). See listed references for otherestimates, including ones with lower Ca, A1, and U. e. Eucrite
parent body. Morgan e t al. (I 978a); bracketed elements estimatedfrom expected cosmic ratios; i 1.2 wt ~ Fe is in the core. See listed
references for other estimates, f. Mercury. See text for simple
model based on oxidized nebula. Minor S, C, and N would be
expected for condensation from reduced nebula, g. Mars. Fe-richmantle proposed by Smyth and McGetchin (I978); note that thismodel would have less Fe if the moment-of-inertia is reduced.
Minor and trace elements not considered.
Note. All the above compositions are based on the theory ofequilibrium condensation with 'stepped' depletion factors for anoxidized nebula with C/O = o.6. Somewhat different compositions would result for a reduced nebula, especially for C, N, S, Cr
and elements which become chalcophilic under reduced conditions.

pipe and by Sibley and Wilband (1977) for average
sediment. For the upper mantle Smith used themean of twenty garnet peridotites calculated by
Carswell (which is very close to a new estimate by
Maaloe and Aoki, I977), together with a rag-bag oftrace-element data culled from many papers. The
Carswell composition was also used for the remainder of the mantle, except that those elements thatshould enter either basaltic liquid or FeS-richliquid during partial melting of peridotite wereassumed arbitrarily to occur only in IO ~o of the
mantle. Ringwood's (1975) model for a lower
mantle that had not lost basaltic component would

M I N E R A L O G Y OF THE P L A N E T S
require a higher content of Ca, A1, etc. than theSmith model; furthermore the Anderson model
would require substantial increase of Fe withdepth. Sulphur was assumed to be the major lightelement of the core, but the original concentrationof 14 wt~o should now be reduced to c.Io wt%(Ahrens, i978), and even lower if other lightelements are present (Si, C, O?). Minor element
concentrations of meteoritic troilite vary considerably with redox state, and Smith assumed that
substantial Cr (I wt%) and other chalcophileelements are present in the FeS component of the
Earth. No single type of iron meteorite was suitable
for modelling the core, and Type III may be slightly
better than Types I and II. Large variations of trace
refractory metals occur in Types II and III irons,
probably because of crystal-liquid fractionation,and log-median values were taken.
Ignoring details, the important features areshown in fig. IO. The ordinate is the ratio ofconcentrations in Earth to 1.5 • CI meteorite, withthe factor 1.5 accounting for the excess of volatileelements (mainly H 2 0 and carbon compounds) inthe meteorite. The abscissa is the column number of
the Periodic Table as arranged by Ganapathy andAnders. Elements that accreted on to the Earthwith the same efficiency from an assumed Ci-type
nebula should plot at the same horizontal level in
fig. Ioa. At first sight there is a lot of scatter, butmuch of it can be interpreted in terms of uncertainty in the estimates of the Earth's composition. There is no need to abandon for now theconvenient assumptions behind the GanapathyAnders type of model, though amendments can be
expected in the future. Nine specific comments
follow: Figs. to(b) and (c) show major fractionations of the crust and upper mantle with respect tothe bulk Earth (fig. Ioa), which correspond to
crystal-liquid partitioning during the segregationof the crust and upper mantle from lower zones.
The pattern for the bulk Earth (fig. ioa) has mostelements within a factor of 2 of the requirements fora Ganapathy-Anders type of model, and there arelikely explanations for the deviant elements; e.g.
old inaccurate analysis for Cd; uncertainty oflocations of Hg and Bi; substantial atmosphericloss of H, as shown by arrow. Magnesium is slightly
higher than Si because of the high olivine content of
periodtite, but could be reduced by putting substantial MgSiO3 (perovskite structure) in the lower
mantle, or by decreasing the Fe/Mg ratio there. The'metal' elements occur at roughly the same level as
Mg and Si, and will tend to become slightly higher if
extra Fe, Cr, and P are placed in the lower mantle
and core. The elements of the 'early condensate'
scatter considerably, but plausible adjustments areavailable: e.g. increase of the Ca and A1 corn-

33

ponents of the lower mantle on the assumption thatit has not differentiated upwards all its 'basaltic'
component (Ringwood, 1975; Tatsumoto, 1978);reduction of Ba by assigning it only to the upper 2
5% of the mantle in conformity with probableoccurrence only in amphibole and mica; adjustment of the rare earths by garnet-pyroxene partitioning (many papers listed in Smith, 1977); increase of V in the core. The 'lithophile volatiles
I3OO-6OO K' scatter widely except for Rb, K, andMn, but F should probably be reduced in order
to concentrate it into the upper mantle, and N aprobably should be increased greatly, as shown by
the arrow, because it enters garnet at high pressure
and might couple with high-valence cations (Si, P,
AI?) in deep-seated minerals; only Cs is likely to beaberrant, and indeed, it is expected to be morevolatile than its companions in the solar nebula.
The 'siderophile volatiles 13oo-6oo K' match quite
well within the large uncertainties. Particularly
important is the contribution of the troilite component to Cu and Zn, but not to Ga and Ge, which
are siderophilic; removal of S and chalcophilicelements from the Earth's core, as in the Ringwood (I977) oxygen-based model, would drastically worsen the fit thereby providing a powerfulargument in favour of substantial S in the Earth's
core. The 'volatiles < 6oo K', shown as open
squares, scatter widely, but there are plausible
reasons for ignoring Cd, B, and Hg, and lowering Pb; probably the most reliable elements are
Th, C1, Br, and I, which concentrate in thecrust and which match quite well. Most of the C
and N is assigned to the core by analogy with ironmeteorites, whereas H can be doubled by assuming atmospheric escape. Because there is a large
tolerance in these elements there is no need toabandon a simple model at this stage, thoughspecial processes may be involved for the very
volatile elements as suggested by many datafor carbonaceous meteorites.
The key role of U as the control element for the
early condensate necessitates detailed discussion. If
all the radioactive heat were instantly transported
to the surface and there were no other heat sources,
the U content of the bulk Earth would be nearo.o3o ppm (there is some uncertainty involved
with processes at the ocean ridges) assuming plausible U/Th and U/K ratios. For CI-proportions of'early condensate', 'silicate' and 'metal', the U content should be c.o.oI 4 ppm (Table IV, column b).
Smith (I977) questioned the validity of estimatesof U content from heat flow, and Daly and Richter (I978) showed that convection processes could
retard heat flow up to about two-fold. By assumingthat Ba has been differentiated entirely into theouter 2oo km it may be possible to estimate its

34

J.V. SMITH

abundance accurately from crustal and uppermantle rocks. In the meantime it seems certain thatthe elements of the 'early condensate' are notdepleted and might be augmented by ~.5(+_o.5?)
over (Mg + Si), as in Table IV, column c.
The bulk Fe content of the Earth has a strong
effect o n the mean density, which is accurately
known. When the equation-of-state has been determined accurately for possible high-pressure phases,
it may prove possible to place accurate bounds on
the ratio of Si to the sum of the other major
elements. In the meantime it is certain that Fe is notdepleted and might be augmented by 1.2(+0. I?)
over (Mg + Si) in CI composition. Actually the Fecontent of C~ meteorites is poorly defined (Kerridge, I977a), and substantial separation of Fe andsilicate has occurred in chondrites (especially enstatite chondrites); hence it is not really certain what
was the Fe/(Si + Mg) ratio of the nebula, and theabove value of 1.2( _+o. I ?) used in Table IV, columnc, must be treated cautiously.
The bulk K and Rb contents of the Earth arehighly controversial, though there is no need to
invoke substantial difference of the K / R b ratio of
the Earth from the C I reference. An absolute lower
limit of9o ppm K results from the assumption thatall 4~
is in the atmosphere, and a modeldependent upper limit of 17o ppm is based on theassumption that all Ar was released to the atmosphere as easily as H20. Unfortunately there arethree permissive factors: present-day emission ofnoble gases from the mantle argues for incomplete
differentiation, and hence for retention of substantial K in deep-seated minerals; pyroxene may bestable down to c.4oo 5oo km, and may carry
substantial K as mentioned earlier; and substantial
K may enter a S-rich core, especially at highpressure (Goettel, 1974). I would not like to increase
the abundance of K more than two-fold in fig. Io(a)because there is no plausible way to increase theabundance of K over that of N a in nebularcondensates and indeed it seems more plausible toreduce the Na content of the mantle with depth
rather than increase it, and secondly, because K

could be flushed outwards with 'basaltic' liquid
during early differentiation of the Earth, and notreact at high pressure with sunken FeS-richmaterial (Smith, i977).
Obviously there are many plausible and nondefinitive arguments which could be discussed
indefinitely, but the following five chemical featuresof the Earth must be explained in any model for its
origin: First, its bulk composition corresponds
generally to accretion models involving only slightfractionation of the 'early condensate', 'silicate',and 'metal' components, and major depletion ofelements expected to condense below I3oo K andespecially below 6oo K. Second, major crystal
liquid differentiation of the Earth flushed out
elements which enter 'basaltic' liquid. Third, substantial volatilization of the bulk Earth is ruled out,
except perhaps at an early stage of growth, becauseof the similarity of depletion factors for elements ofwidely different thermal volatility [note that'thermal volatility' is quite different from volatility
in the solar nebula]. Fourth, siderophile andchalcophile elements were retained substantially inthe upper mantle and crust (fig. Iob, c) at a late stageof accretion, and fifth, storage of major amounts of
Ca, A1, Na, Ti, etc. occurs deeper than 2oo km.
Furthermore, there is the probability of substantial
loss of H through the atmosphere, and major
storage of C and N in the core and mantle.
Detailed discussion is needed on the variation ofredox state with depth. Certainly the present crust
contains substantial ferric iron, while upper-mantle
peridotites have a high ratio of Fe2+/Fe 3+. The
occurrence of diamond in peridotites, and theexpulsion of CO2 in carbonatitic and kimberlitic
magmas, together with the indication that volcanic
gases are buffered b y magnetite silicate (Nordlie,t977), testify to absence of extreme redox states in
the upper mantle. However, it is not clear what is
the redox state of the lower mantle and core,especially in view of possible effects of very highpressure on transition metals and even alkali
metals. It is probably wise not to rule out thepossibilities that the Earth accreted substantial

FIG. Io. Relation between Earth composition and CI composition. The base is labelled with column numbers from the
Periodic Table, and the elements are classified according to the sequence of equilibrium condensation of a solar nebulawith C/O c.o.6 and I o Pa pressure. The ~.5 factor allows for removal of very volatile elements from C I composition. (a,top) Model III for the bulk Earth is based on 38 wt % troilite of meteoritic composition and 62 wt % type IIIAB ironmeteorite in the core and garnet peridotite composition for the entire mantle except for occurrence of large-ionlithophile elements only in the outer IO% (putting Na in the whole mantle would raise Na above Rb and K as shown byarrow), 'andesitic' crust, and the hydrosphere and atmosphere. The primordial content of H might have been about
three times higher as shown by the arrow. Particularly uncertain elements are shown by question marks. Ranges forelements result from element partitioning in type III meteorites, and the symbol is placed at the log median value. (b,middle) Earth's crust. (c, bottom) Peridotite (mainly garnet) as model for uppermost mantle. The significance of this typeof plot is explained in Ganapathy and Anders (i974), and the many sources of data, along with detailed tables, are given
in Smith (1977).

,~Lu

,.'~

ne

~ j

l.I.

i 9 Early condil

9 il K F z _ _ _ _ _ _ _ q l M .

co--

IO-'J-"

u

,,,,

'

i~

us

8

.

102 I BoaK

OLa OTh 9t ZrlH f To

9

IIRb
--QLi
iCs
iNo 9

slO
E

W,U
9

[] ,?n
DTI

Br?F1

[o0

ci?D
~ N

T

H~]??

8

1B 2B 3A 4A 5A 6A 7A
I

I

v
i

I

I

I

I

.Mo

Mn
9

OH

Co

Cr

9

I

I

I

I

[]Ge

~ [E

K

I

I

I

I

I

I

i i I i I ?

I

i

OTh
Sr

IC s o"c-"~ 10-1 --

~

IHf

-

i

i

i

i

i

0 [] Volaliles<600 K/ . . . . . . . . . . . . . . . . . .

W
Or

LUly
..............

"~

9
Mn 9

9 EarlyCondensate ',Silicate

9

OH

1

Lo

ill~-wLr-- 9149
No

!
I
I

QU

Bo

S
[]
Se

1B 2B 3A 4A 5A 6A 7A
9Silicole
9
Metal,D
Volaliles<600 K

1A 2_A 3B 4B 5B 6B 7B
8
9 Early Condensole
9 Lithophile l Volatiles[] Siderophile]1300-600 K

tO 2

N
0

I

B .

Fqln

Au

Ni

I

I

Pb

OMg

I

I

TI [] Sb[] Sill []
IIF
BC]
C~g
" []~-[]Sfi-&Bi~lBrCd
P As 0
~Zn
lC
9

Fe

I

A

OAI

9

1(2

Te

z.

~1
Y, Lo 99Lq,"~ lo-1

Pb?

[]

It

1A 2A 3B 48 5B 6B 7B

Lmlr

i ~#[]

Flsn'/''

I ~o .i,I Silicate
II) / I
9imcs iI " . . . . .MetalI
N~ I
V01atiles!
/
10-2_[ 1300-600 K i
I/J~, i 9 Lithophile[] Siderophlle L,___ q
I oH I [] Volatiles < 600 ~,
I

_
Et

[]

',

Metal

Volatiles

Fe

Pd9

I

13oo-6OOK ,,,9 LithophileI[] Siderophile

I

q
J i
f J',r
1A 2A 3B 4B 5B CoB 7B

9Cu

r~Zn C]InP

HI

apt ~A0?

J
I Re

Ni

Si

i

B~Ti ~ !ASBi ~ i F
r~Go []Sn
clmBr

9
9
Co

AI

Ir

9

~Ge
~IS

~ 9

~

i

8

C

N

0

0

[]Se

I I I I I I lIB 2B 3A 4A 5A 6A 7AC36

J. V. SMITH

amounts of highly reduced material, and that theouter 2oo km became oxidized either from lateaccretion of oxidized material or from major
atmospheric loss of H and subsequent subductionof oxidized rocks.
It is not clear how much of the radial heterogeneity of the Earth is the result of heterogeneous
accretion, and how much is the result of postaccretion crystal-liquid differentiation. At one
extreme, accretion might have been near-homogeneouswithcrystal-liquid differentiation
proceeding by partial melting, whereas at the otherextreme, accretion might have been strongly
heterogeneous with 'early condensate' and 'metal'
fractions accreting first. Whatever the outcome of

discussion, cosmochemical modelling implies thatthe middle and lower mantle contain large amounts
of Na, A1, and Ca, and that the early differentiation
swept out only some of the 'basaltic' component tothe surface. Further discussion is needed on theimplications of this important conclusion, but it
may mean that the Earth was never fully molten, orthat it was not molten after it had accreted to (say)half its present mass. Perhaps the Earth is now
moving increasingly from an early era dominatedby crystal liquid differentiation of 'wet' K-bearing
peridotite to a new era dominated by second-stage
differentiation of deep-seated 'dry' K-free peridorite. Will plate-tectonic processes involve subduction to increasingly greater depths?

MOON'Then... the moon must be older than the earth?'. 'No!' said Barbicane
decidedly, 'but a world which has grown old quicker...'Jules Verne, From the Earth to the Moon
This section is deliberately brief because of theexistence of numerous overlapping reviews (e.g.
over-all view, Taylor, i975; classical mineralogy,
Frondel, 1975; petrologically oriented mineralogy,
Smith and Steele, 1975; survey of lunar rock types,
Wood, I977; mare basalts, Papike et al., 1976;
Papike and Vaniman, 1978; ANT-suite rocks, Prinz
and Keil, 1977; KREEP-rich basalts, Meyer, 1977).
For brevity, detailed references are given here only
to controversial subjects. Readers can find appropriate up-to-date primary references to most subjects in the Proceedings of the Lunar Science
Conferences, and secondary reviews in many places,
including Massey et al. (I977), Reviews Geophysics
and Space Physics, and Phys. Chem. Earth.
Observations

Although the Moon's surface has been reworked
by impact of bodies ranging from tens ofkilometresdiameter down to micrometeorite size, the absenceof plate-tectonic processes and the lack of mineralalteration has allowed reconstruction of the maj or
events in the Moon's petrologic history merely
from detailed study of surface rubble returned fromnine sites. Most returned homogeneous rocks aremare basalts attributed to late volcanic flows in
impact basins, and the early lunar crust is represented mainly by breccias, though a few homogeneous rocks survived (Phinney et al., I977).
Painstaking study of tiny fragments in breccias andsoils was one of the keys to reconstruction of thepetrologic history of the Moon, while photogeologic mapping coupled with geochemical mapping from Earth-based telescopes and orbiting
satellites was essential for placing the petrologic

information in a geochemical and geophysicalcontext. Great emphasis was placed on thechemistry of minerals and rocks in the developmentof petrogenetic models, and the lack of stratigraphic control was thereby mitigated.
The lunar rocks and minerals can be crudelyclassified into five groups: Very rare 'ultrabasic'
material (mostly small fragments and mineral
grains) contains Mg-rich olivine and pyroxene,
which can be attributed to early products ofcrystal-liquid differentiation. A second group ofrocks, minerals, and glasses has chemical, andsometimes petrographic, properties similar toanorthositic, noritic, or troctolitic rocks: this ANT
suite is attributed to an early heterogeneous crustrich in plagioclase, low-Ca pyroxene, and olivine. A
highly diverse group of fragments (mostly frommetamorphosed
breccias), whose
acronym
KREEP
referstoenrichmentin
K,
REE(Y, Ln), P, Ba, Th, and U has a considerable
range of enrichment and bulk composition;although separation of 'basaltic' liquids fromolivine, plagioclase, and pyroxene crystal residues
is needed to give the enrichment of this third group,
it is not clear what are the relative contributions ofprimary crystal-liquid differentiation, of secondary
partial melting of crystalline cumulates withtrapped liquids, and of multi-step processes andhybridism; particularly confusing is the controversy over glasses and heterogeneous rocks with
Low-K-Fra-Mauro basaltic composition, whichcontains substantial KREEP component, butwhich may have a variety of origins (Reid et al.,1977). A fourth group of mare basalts and mafic
glasses with wide variation of Fe, Ti, and A1 hasgenerally high ratios of Fe/Mg, which requireM I N E R A L O G Y OF THE PLANETS
advanced stages of crystal-liquid differentiation, as
do the trace-element signatures; partial melting ofvarious heterogeneous Fe-rich cumulates (somealso Ti-rich or Al-rich) is the most popular mechanism for origin of mare basalts (e.g. Shih andSchonfeld, 1976), but the original sources of the
cumulates, the extent of crystal-liquid fractionation during emplacement, and the possibility ofhybridism of magmas are the subject of intense,
non-definitive debate. Finally the debris frommeteoritic bombardment is represented by rarefragments of meteorites, metals, and alloys, and by
ghostly signatures of non-lunar trace elements
attributed to dispersed meteoritic debris.
Numerous experimental syntheses at controlledtemperature, pressure, and oxidation state provide
controls on the possible crystal liquid fractionations in the Moon, but there is considerable
controversy over the depth of origin of marebasalts because of disputes over the crystallization
sequence near the solidus, and of uncertainties
about the effect of crystal separation on the bulk
composition of supposed primary magmas. Certain
elemental ratios (e.g. Mg/Fe and Cr/A1/Ti) provide
good guides to the direction of differentiation.
Many lunar rocks can be represented closely bycompositions in the olivine plagioclase-silica ternary system, and there seems to be a growing
consensus that olivine dominates over pyroxene in
the bulk Moon; furthermore, the pyroxenes tend
to be Ca-poor except in mare basalts. Numerous
trace-element ratios provide fairly good controls
over the efficiency of crystal-liquid differentiationon the Moon, and are the basis of several estimatesof the bulk composition.
Geophysical data provide eight constraints:
First, the density of 3-34 g/cm3 enforces an upper
limit of c.4oo km for an Fe-rich core and c.7oo km
for an FeS-rich core if the mantle is dominated by
Mg-rich olivine (atomic Mg/Fe c. IO). Second, estimates of the moment-of-inertia, which have fallen
by amounts greater than stated errors to recent
values near o.392, require a centre-heavy Moon.
Third, seismic profiles (Goins et al., I977), whose
interpretation is complicated by closely similartransmission properties of Ca-rich plagioclase andMg-rich pyroxene are consistent with a crust rich inplagioclase and pyroxene overlying a mantle rich inolivine; furthermore, poor transmission of S-waves
through the Moon's centre requires a hot interior.
Fourth, the offset of 2 km between the centres ofmass and figure, when taken together with otherdata, requires a light plagioclase-rich crust, ranging
in thickness from perhaps 5o_+ 25? km on the nearside to lOO_+5o? km on the far side, lying on adenser mafic mantle. Fifth, remanent magnetism ofearly lunar rocks, especially breccias, is controver-37

sially attributed either to a dynamo in an earlyliquid core or to special processes involving external magnetic fields that were especially effective
during impact brecciation. Sixth, the electricalconductivity profile, derived from the perturbationof the solar wind, yielded model-dependent temperature profiles, which allow, but do not require,
crystal-liquid differentiation of a lunar interior of
Mg-rich silicates. Seventh, estimates of the radioactive heating from the surface heat-flow indicate,but perhaps do not definitively require, two-foldenrichment of U in the bulk Moon with respect todevolatilized Ci-meteorite composition, and,eighth, the occurrence of mass anomalies only forlate basins shows that such mare-basalt fill arrived
too late for isostatic adjustment with adjacent nonmare regions.
Orbital missions have provided detailed photogeologic maps, which were compared with chemicalmaps of some of the equatorial regions for
Mg/A1/Si (X-ray technique) and Th, K, U, and Fe(gamma-ray technique). Important results are: afirst-order distinction between mare and non-mare(highland) regions; second-order distinctions
correlatable with details of mare flows and impact
ejecta; a general correlation between crustal thickness (estimated from isostatic compensation ofelevation) and chemistry (Fe negative; A1 positive;
Th complex); irregular concentrations in theImbrium-Procellarum region of KREEP-richmaterial; and a tendency for A1/Si/Mg to changewith longitude for both inter-mare and highlandareas so that the feldspar content appears toincrease from west to east on the near side. Spectralreflectance data from earth-based telescopes werecorrelated with direct observations of returned
samples to show that several types of mare basalt
occur in the same basin, that several mare units are
not represented in returned samples, and thatnumerous dark deposits of possible pyroclastic
origin need to be sampled.

Petrologic interpretations
Taken together, these data provide a goodframework for a general petrologic model of theMoon, but the details are highly tentative anddifficult to test definitively because of the considerable tolerance in estimates of relative proportionsof rock types. The following model incorporates
chronological information obtained from isotope
studies.
First, the Moon underwent major crystal-liquid
differentiation at c . - 4-5 x I O 9 yr, probably involving complex diapiric processes triggered by impactof large bodies, and perhaps involving the entire
Moon. The differentiation produced lbur products:

38

J . V . SMITH

a heterogeneous 'crust' composed of feldspar-richrocks (including anorthositic, noritic, and troctolitic varieties ranging from plutonic to volcanic
types); a heterogeneous 'mantle', perhaps statistically stratified from a lower mantle rich in olivine
and Mg-rich complex oxides to an upper mantle
rich in low-Ca pyroxene and complex oxides lessrich in Mg (note that simple crystal-liquid separation of a completely molten Moon could haveyielded a gravitationally unstable mantle withheavier Fe-rich minerals lying above lighter Mgrich minerals, whereas convective overturn or local
diapiric processes could have yielded different
spatial distributions); residual liquids, rich in
KREEPBaThU
components, which tended
statistically to concentrate between the 'crust' and'upper mantle', but which might have become
distributed erratically throughout the 'crust' and'upper mantle' by volcanic and impact-induced
processes; and perhaps a small core of Fe,S-richmaterial9 The time-temperature-depth history of
the Moon is poorly constrained in spite of numerous model calculations, and whether a single deep
'magma ocean' actually existed is highly speculative. The extent of gravity-separation or elutriation
is debatable, and plagioclase flotation in an asymmetric gravity field, controlled by a near-by Earth,is just one potential way to develop the Moon's
asymmetric crust, while differential bombardment
is another.
Second, the primary differentiation tailed off into
a period of complex secondary differentiation
whose details are quite uncertain9 Perhaps little
happened in the 'lower mantle' and 'core', while themajor action occurred in the 'crust' and 'upper
mantle' which were subjected to impacts from lateplanetesimals, and to radioactive heating, especially from KREEPBaThU-rich regions. Partial
melting in the 'crust' may have resulted in cotectic
melts, while huge impacts may have homogenized
portions of the crust giving a range of rock types
and Low-K-Fra-Mauro compositions. Whether
there was a major basin-forming 'cataclysm' at- 3.9 x IO9 yr or whether the recognizable basins
were formed randomly f r o m - 4 . 3 t o - 3 . 9 x i o 9 yr
is highly controversial, but from the petrologic
viewpoint there can be no doubt that much of thelunar surface was affected by complex metamorphic processes associated with major impacts(Phinney et al., t977). An alternative to the suggestion of a quiescent 'lower mantle' is that trappedliquid from the primary differentiation seeped
upwards and hybridized with the material in the'upper mantle' and 'crust'. Whether the lunar crustis stratified into upper 'anorthositic gabbro'
composition overlying middle 'Low-K-FraMauro' composition and lower 'mafic ANT'

composition (Ryder and Wood, 1977) is debatable,
as are the relative effects of primary and secondary
differentiation in achieving such a presumed
stratification.
Third, cessation of the basin-forming impacts,
and reduction of the radioactive heating, allowed
the lunar surface to largely stabilize into the present
stratigraphic provinces of ancient highlands,
impact debris, old irregular maria, and young
circular maria. Partial melts from the 'upper
mantle' yielded mare basalts. These may have been
forming before cessation of the basin-forming
impacts, and the returned mare basalts may beproviding information mainly on the final cappingof the maria after a presumed long history of
igneous and metamorphic activity in the early
impact debris and the uprising mantle (and crust?)
under the impact basin. The late mare basalts musthave arisen from considerable depth in the Moon
(c9km?) in order to allow a sufficiently rigid
crust above the weak source region, but the estimates of depth from presumed cotectic birth of themagmas is controversially dependent on subtle
details of experimental phase equilibria9 Crystallization ages of mare basalts range f r o m - 3 . 9 to
- 3. I x io9yr, but earlier ones may have been lost in
the metamorphosed debris of impact basins, whilelater ones indicated by photogeologic mapping did
not turn up in samples returned to Earth9 Largecompositional variations o f F e , A1, and Ti in the
mare-basalts and mafic glasses indicate derivationfrom heterogeneous source, regions, or crystal
liquid differentiation during:ascent and emplacement, or both. Paucity of mare basalts on the far
side of the Moon might be explained by the greaterhydrostatic head needed to penetrate the thicker
far-side crust, or by lower production of marebasalt in the far-side mantle, or both.
This generalized model stitches together andmodifies the original general ideas of early majordifferentiation to produce a feldspar-rich crust andan olivine-rich mantle (Smith et al., 197oa, b; Woodet al., I97O) and partial melting of igneous cumulates to produce mare basalts (Philpotts andSchnetzler, I97O). Probably most lunar petrologists
are currently accepting these two general features,but many of them prefer to differentiate only some
hundreds of kilometres rather than the entire
Moon. Detailed discussions of crystal liquid differentiation in a magma ocean, together with controls
imposed by mineral and rock chemistry (e.g. Mg/Fe
ratio, (Y,Ln) signature, Ni and Cr contents) can befound in recent Proceedings of the Lunar Sciences
ConJerence.
A second type of model was developed from theoriginal proposition by Ringwood and Essene(I97O) that Apollo I i basalts originated by partialM I N E R A L O G Y OF THE P L A N E T S
melting of a (primordial) pyroxene-rich Moon
generated from a hot fission cloud thrown off the
Earth. The latest version (Ringwood and Kesson,
1977) has adopted the Smith-Wood concept of anearly major differentiation of the Moon (but only
to 4o0 km depth), which produced a gabbroic crust(c.5o km thick) overlying an intermediate layer(c.5o km) and a refractory residue of dunite(c.3oo km), but invoked generation of mare basalts
by partial melting of a diapir rising out of deepseated (c.5oo km) primordial olivine pyroxenite
into the overlying dunite zone. Presumably thecentral zone of the Moon still remains as undifferentiated primordial pyroxenite.
Untbrtunately it will prove difficult to distinguish the Philpotts Smith Wood class ofmodels from the Ringwood-Kesson model becauseof four reasons. The RK model has adopted anearly differentiation of half the Moon's volume
according to the Smith-Wood model; the rising
melt from the RK diapir can interact with the
cumulates produced in the early differentiation,
which are the source of mare basalts in the
Philpotts model; the experimental partition coefficients for element distribution between liquid andcrystals have considerable tolerance; and the initialchemical parameters of model Moons (e.g. Mg/Si
ratio) have some tolerance. Geophysical tests areprobably inconclusive because of poor control overthe properties of the Moon's interior. If the lunar
magnetism could be uniquely attributed to a
dynamo (Runcorn, i976), complete differentiationof the Moon would be indicated. Ringwood andKesson (1977) pointed to the problem of heatsources for melting the entire Moon, but Toks6z
and Johnston (1977) were able to develop a thermal
model for complete differentiation of the Moon.
Brett (1977) also questioned whether the bulk of theMoon melted. Solomon and Chaiken (1976) suggested that the absence of major fractures in the
lunar surface limited an early 'magma ocean' to
only 2oo_+ IOO km depth (now revised to 3oo_+ IOO
km; Solomon, pers. comm.), but detailed study isneeded to check whether such a thin ocean can
provide a plagioclase-rich crust thick enough(c.75 kin?, Kaula, I977a ) to explain the offset between the centres of figure and mass and allow sufficient trace-element enrichment of mare basalts.
Probably it will be difficult for geochemists,
mineralogists, and petrologists to come up withconvincing controls on the deep interior of theMoon, but there is great scope for further study ofnear-surface differentiation.
Bulk chemical composition

Fortunately from the viewpoint of planetary
mineralogy there is a substantial convergence of

39

opinion about the bulk chemical composition of
the Moon. This arises mainly from the tendency toassume that at least half of the Moon's mass
participated in an early major differentiation into aplagioclase-rich crust and an olivine-rich complement, together with likelihood that the undifferentiated interior (if any exists) is dominated by
Mg-rich silicate rather than non-peridotitic compositions. Early extreme views, such as a Moon
dominated by Ca,Al-rich nebular condensate(Anderson, I973) , are quite unacceptable when thesurface rocks are dominated by 'basaltic' compositions, and recognition of the uncertainty of theestimate of U content from heat-flow measurements (Langseth et al., 1976) has allowed reductionof early estimates of high U and cosmochemicallyassociated Ca, A1, and Ti (Ganapathy and Anders,I974). All recent estimates of lunar composition(Taylor, 1976; Binder and Voss, I978; Taylor andJakeg, 1977; Dreibus et al., 1977; Kaula, 1977a;
Kesson and Ringwood, 1977; Longhi, I977;
Ringwood and Kesson, I977) are dominatedmodally by olivine, pyroxene, and plagioclase,
as in early models by Smith et al. (I97oa, b)
and Wood et al. (197o), and can be described as
peridotitic.
The eight key features of all recent model Moons(Table IV) are: the low content o f F e O (IO-I 4 Wt~o),whose maximum value is strongly constrained by
the mean density for all models dominated by Mgrich olivine; the very low content of volatileelements, as typified by the calcic nature of almost
all lunar plagioclase (thereby requiring low N a andK), and the almost complete absence of H 2 0 inlunar rocks; the very low content of metal-seeking
elements (e.g. Ir and Au) in lunar surface rocks,
thereby allowing the extent of meteoritic contamination to be estimated by measuring theseelements in soils and breccias (Hertogen et al.,I977); the dominance of Mg-rich olivine in the
differentiated lunar interior in order to allow
primary crystal-liquid differentiation to proceed to
ANT-like compositions (e.g. Drake, I976; Longhi,
I977); the limitation of Mg/Fe ratio in the olivines
by observed mineral compositions and crystalliquid fractionation factors; the limitation of Ca
and A1 from constraints in the amount of Ca-richpyroxene and plagioclase produced during differentiation (recent estimates ranging from 4 to 8 wt
of AlzO3 are lower than earlier values of 1I wt
linked to high estimates of U from the heat flow);
the limitation of many trace elements from model
calculations of crystal liquid differentiation (e.g.
Taylor, I975); and absence of precise control over
the amount of S in a hypothetical lunar core, exceptthat an upper limit of I w t ~/o seems likely if theMoon is dominated by Mg-rich olivine.

40

J . V . SMITH
ORIGIN OF THE EARTH AND MOON'... and who can say that the moon has always been a satellite of the earth?'
Jules Verne, From the Earth to the Moon

Because the origin of the Earth and Moon is sohighly controversial, the existing geochemicalmineralogic-petrologic data must be utilized asmuch as possible in placing boundaries on models.
If one accepts the concept of planetesimals asintermediaries between nebular condensates andplanets, one must expect the Earth-Moon system
to have developed in a complex way. Perhaps most
lunar scientists would now agree that the three
simple models of fission of the Moon from the
Earth, non-disintegrative capture of the Moon by
the Earth, and simultaneous accretion of the Moonand the Earth, are insufficient, and that it isnecessary to move to more complex models.
Readers should note that I have discussed disintegrative capture of a large proto-Moon (Smith,I974), but in this review I shall try to stand aloof in
the tradition of Chamberlin's Method of Multiple
Working Hypotheses, and indeed now suggest thatno single process can explain the origin of the
Earth-Moon system.
Age constraints
The crystallization ages of the mare basalts
require that the Moon was orbiting the Earth atleast by - 4 . o x I O 9 yr, and it is easiest to interpret the model ages of radioactive species in termsof an early major differentiation at c . - 4 . 5 • IO9
yr. Capture of the Moon by the Earth would beso catastrophic that major geologic provinces
would have been obliterated. It is easiest to
assume that the lunar chronology began with theMoon in Earth orbit by - 4 . 5 x lO9 yr.
The Moon was certainly bombarded by large
projectiles up to c . - 3 . 9 x io 9 yr, and the Earthmust have been affected by the same populationof bodies. Furthermore the effects on the Earthshould have been more catastrophic than on theMoon, and it is necessary to consider impacttriggered volcanic processes in the outer part of the
Earth up to c. - 3-9 x io 9 yr. This is consistent withthe earliest crystallization ages of c.-3.8 x IO9 yr
for terrestrial rocks.
Chemical constraints
This section is based largely on Smith (1977,1978).
Incomplete equilibration of the Earth is indicated by the well-known occurrence of substantial
amounts of metal-seeking elements in the Earth's
upper mantle (e.g. 0. 3 wt ~o Ni in upper-mantle

olivine). No matter whether the Earth accreted
homogeneously or near-homogeneously, it isnecessary to assume that a mantle barrier blocked
equilibration between the upper mantle and core.
Unfortunately it is extremely difficult to calculate
the rate at which equilibration would occur even in
a completely molten Earth, but it is qualitatively
easier to envisage lack of equilibration in an Earthcool enough to have a solid mantle when (say) 8o9o ~o of the Earth had accreted.
The alkali metals in the bulk Earth are depleted
not more than six-fold with respect to water-free,
Ci-meteorite composition, and it is difficult or
impossible to retain these alkalis in Ringwood's(i975) volatilization model in which refractoryelements such as Mg, Si, Ca, and A1 are supposed tohave volatilized from a very hot Earth to form theMoon.
Estimates of the bulk composition of the Earth
are subject to considerable uncertainty from lack of
direct knowledge of the lower mantle and core. It is
possible to satisfy geophysical constraints withmodels in which the major elements Si, Mg, Fe, Ca,
and A1 have nearly the same ratios as in CI
meteorites, though perhaps there is enrichment of
Ca and A1 (I.5-fold?) and Fe (i.2-fold?). Volatileelements are considerably depleted, but there is noneed to attribute the depletions to any special
factors in the origin of the Earth-Moon system;
rather, it is initially easier to attribute them toprocesses involved in nebular condensation (Ganapathy and Anders, I974) and catastrophic accretionof planetesimals (many authors).
Perhaps the chemical composition of the Moonis more closely constrained than that of the Earth.
It has several important features (Kaula, 1977a):
low abundance of Fe (three-fold lower than
Earth?), and almost certainly of all siderophileelements; not more than two-fold enrichment ofearly-condensate lithophile elements with respect
to C 1 composition (Ca, A1, U); very low abundanceof all volatile elements with respect to CI composition, and even with respect to Earth; and dominance by Mg-rich olivine composition, as is also
likely in the bulk Earth.
Dynamical and accretional constraints
Origin of the Moon by simple rotational fission
faces momentum and energy problems (e.g. Kaula,
1971, i977b), but these can be obviated by combining rotational fission with other mechanisms.

M I N E R A L O G Y OF THE PLANETS
It may be possible to develop useful constraints
on accretion. If Safronov's models for the size
distribution of accreting planetesimals can be adequately developed to cover the complex processes
involved in collisions, it may prove possible to place
an upper limit on the size of planetesimals; current
theory is incompletely developed (Wetherill, 1976).
Although the ratio of escape velocity to impact
velocity is extremely important in determining thefate of colliding bodies, it is too early to rely onmodel calculations; present calculations suggestthat the Moon is and was very much less efficient inretaining collision debris than the Earth (e.g.
O'Keefe and Ahrens, I977a , b), especially formaterial arriving at IO km/s or greater.

Comparison of suggested origins
Fission. Shoemaker (I977b) suggested that accretion of the Earth from large planetesimals would
inevitably result in major changes of angularmomentum as the result of random off-centre
collisions, and indeed he proposed that off-centre
collision by one or more particularly large
planetesimals would lead to one or more episodesof fission. Actually one can expect a whole gamut of
events ranging from head-on collision via glancing
collision (Hartmann and Davis, I975; Cameronand Ward, 1976 ) to disintegrative capture ofmechanically weak bodies in the Roche instability
zone (Opik, ~972; Smith, i974; Mitler, i975) , and it
may prove impossible to estimate the relative
proportion of the types of events. Three types ofchemical differentiation can be envisaged: on
Earth, such that a piece of the mantle was flung or
torn off, leaving the lower mantle and core in place(Binder, ~976; Ringwood and Kesson, 1977; Rammensee and W~inke, i977); in orbit, such thatvolatiles were lost, and mechanical differentiation
occurred for fragments of different chemicalcomposition (Smith, i974; Kaula, ~977a, b); andthirdly, during partial capture of the colliding body,which initiated fission (Smith, 1974; Mitler, 1975),
especially if it had already differentiated chemicallyin heliocentric orbit.
There is no great difficulty in accounting for a
Moon low in volatiles by a fission-dominatedevent. There is no problem in obtaining an olivinerich Moon since both the outer mantle of the protoEarth and colliding planetesimais might be olivinerich. Deficiency of Fe and metal-seeking elementsin the Moon could be explained by obtaining theMoon largely from the outer mantle of the protoEarth; furthermore, the deficiency might be augmented by more efficient capture of silicate debris
from planetesimal(s) that initiated fission. Smith(1977) emphasized that it is difficult to retain

4I

substantial amounts of volatile elements on Earthif simple rotational fission had occurred, but in1978 pointed out that late accretion of volatile-richmaterial by the Earth after fission had occurred
could account for the volatile elements so long as
the accreting material arrived at too high a velocity(cAo-2o km/s?) to be retained significantly on the
Moon.
Disintegrative capture. One extreme version ofthis mechanism involves a single large planetesimal(e.4 x present lunar size?) which had differentiatedchemically into a metal-rich core and silicate-richmantle, while another consists of fortuitous headon collision of two planetesimals near the Earth(Levin and Mayeva, I977). The Moon would beformed from debris produced during the capture,and numerous possibilities exist for separation ofsilicate from metal components as in the preceding
paragraph. Less extreme versions might involve
disintegrative capture of many planetesimals,
which produced debris that accumulated over a
period of time to form the Moon, and indeed suchversions might be classified under simultaneous
accretion. Unfortunately it is impossible to pin
down the many possible chemical fractionations
that are available, especially because of the unknown factors involved in the relative mechanicalstability of an iron-rich core and a silicate-richmantle. Many complex possibilities are also
involved in the accretion of Earth-orbiting debris(Opik, I972; Safronov, i972; Kaula, I977b).
From the general similarity of the chemicalcomposition of the Moon to that of the supposed
eucrite planetesimal, Anders (1977) suggested that a
special process is not needed to explain the origin of
the Moon. Indeed it might be argued that a
supposed proto-Moon already had a Moon-likecomposition prior to disintegrative capture. However, the eucrites are particularly depleted in volatile and in siderophile elements, and it might beunwise to press analogies with the Moon.
Since planetesimals growing at the edges ofplanetary feeding zones should grow longer thanones in the middle of the zones, might a protoMoon have formed approximately mid-way between Earth and Mars, and been captured near theend of Earth's accretion?
Simultaneous accretion. This process could nothave begun unless a proto-Earth and orbiting
proto-Moon had formed by either capture (disintegrative?) or fission. Furthermore, simultaneous
accretion must have occurred to a greater or lesser
degree for models dominated by fission or disintegrative capture. Whether the difference of
impact velocity for the Earth and Moon could
account for most of the chemical differences
between the Earth and Moon is debatable.

42

J.V.

Ganapathy and Anders 0974, P- I2OI) argued
that the different chemical compositions of the
Earth and Moon result merely from selective
accretion of aerodynamically sorted dust and thatthe escape velocity is unimportant. However,this ignores accretion of large bodies for which
O'Keefe and Ahrens (I977a, b) imply poorer collection efficiency for bodies impacting the Moon
than those hitting the Earth.
Particularly important are the ideas of Ruskol(I977) for simultaneous accretion of the Earth andMoon, in which great importance is attached to theformation of a swarm of small bodies around theEarth during its growth. The capture efficiency ofthe swarm is greatly reduced if the swarm accretes
into a proto-Moon or several moonlets. Incoming
bodies might disrupt the proto-Moon or moonlets.
Ruskol suggested that the Moon accreted most ofits material only at the end of the accretion periodof the Earth, and that its distinctive chemistryarises because the late heliocentric bodies had
undergone more collisions than the ones captured
early by the growing Earth; these late bodies would
be volatile-poor and richer in brittle silicate than

SMITH

coherent iron. Kaula (I977b) adopted Ruskol's
general model, but augmented it by further discussion of chemical differentiation of impact debris in
Earth orbit. Complex dynamical factors are also
important, and one can envisage complex possibilities for accretion near the Roche limit, andsequential capture ofmoonlets (Opik, I972; Smith,1974).
Combined processes. Currently it seems plausible
to attempt to explain the origin of the Earth andMoon in terms of combined processes. In general, it
seems best to consider fission (coupled with disintegrative capture) only for the first half of the
Earth-Moon accretional history, while simultaneous accretion (again coupled with disintegrative capture) became the dominant process late in
the history. Probably the evidence on the details of
accretion of the Earth and Moon has mostly
disappeared leaving only ghostly hints such as themare basins, and the chemical disequilibrium of the
Earth. The iron depletion of the Moon could be thesum of effects occurring in helio- and geo-orbiting
bodies, and the low abundance of volatiles could be
the sum of several processes.

MERCURY
Look here u p o n . . , the herald Mercury
New-lighted on a heaven-kissing hill.
William Shakespeare, Hamlet, nL iv. 53.
Because of the paucity of direct evidence on
mineralogy, petrology, and chemistry of Mercury,
all ideas ultimately depend on cosmochemical
models. The following account is based mainly on
papers in vol. 28, no. 4 of Icarus.
Mercury has an eccentric orbit (the result of alate large impact?), and a rotation period (58.65
days) precisely in two-thirds resonance with its
orbital period (almost certainly from tidal interaction with the Sun). Its uncompressed density(c.5. 3 g/cm 3) can be explained only by a highcontent of Fe, as correctly adduced by Urey (I 952).
For an Fe,Ni core and a diopside-rich mantle,
approximately two-thirds of the mass is containedin the core. Siegfried and Solomon (1974) estimatedo.66 assuming Feo.9Nio. a in the core, and silicateswith uncompressed density 3.2 g/cm 3 in the mantle.
Substitution of a light element in the core (e.g. S orSi) would require compensating reduction in the
proportion of silicate, and indeed the uncompressed density of troilite (4.8 g/cm 3) is well belowthat of Mercury. The Mercurian atmosphere can be
explained by temporary capture of light species
from the solar wind, and there is no evidence for
substantial volcanic outgassing (Kumar, I976) orsignificant weathering of craters.

Earth-based telescopic study shows a continuous
increase of reflectance from 0.32 to I.o 5/~m similar
to that for spectra of the lunar surface except for
the lack of a significant absorption band at c.o.95#m in the Mercurian spectrum (Vilas and McCord,
I976; Hapke, I977). This absence indicates < 6
wt ~o FeO at the Mercurian surface, and the generaltrend between reflectance and wavelength indicatesa mature regolith, probably dominated by breccias
and glasses. Adams and McCord (I976) suggested
that the surface composition is anorthositic, but I
suggest that other minerals with a featurelessspectrum should be considered, especially as Mercury need not have accreted substantial amounts offeldspar according to simple condensation modelsof the solar nebula: likely candidates are diopside,
spinel, forsterite, enstatite, and melilite. The
Mariner photographs of the Mercurian surface(Murray et al., I975) revealed smaller variations ofalbedo and colour than on the Moon, indicative oflower range and absolute level of Fe and Ti.
The surface of Mercury is heavily cratered, anddisplays conspicuous basins (especially Caloris)
and fairly smooth areas subdivided into pre-basin'inter-crater plains' and post-basin 'smooth plains'.
Long, low cliffs ('lobate scarps') were attributed toM I N E R A L O G Y O F THE P L A N E T S
compression of the crust as a result of continued
cooling of the interior after consolidation of thecrust. Although there is no unequivocal evidence of
volcanism, such as flow fronts, the plains arecommonly ascribed to volcanic activity (Trask andStrom, ~976) though Wilhelms (i976) attributed
them to impact melt or other basin ejecta by
analogy with light plains on the Moon. Lowman(I976) attempted to match the inter-crater plainswith the pre-Imbrian lunar highlands and thesmooth plains with lunar maria, thereby requiring
two episodes of igneous differentiation on Mercury. Solomon (2977b) argued that compressional
tectonic features require early melting. Whatever
the outcome of the speculation on the surface
history of Mercury it seems safe to conclude thatany major crystal-liquid fractionation of the planet
must have taken place before the basin-forming
events (? before - 3 . 9 Gyr).
Particularly important is the firm evidence for a
dipole magnetic field, with its implications on thethermal evolution. Ness et al. (2976) demonstrated
severe difficulties in attributing the dipole to fossil
remanent magnetization, principally because theexpected heat sources should melt Mercury, andbecause subsequent cooling would allow only athin outer shell (probably very low in Fe-metal) to
go below the Curie point. Ironically, the alternative
assumption of a dynamo raises the problem offinding an energy source to drive it at the present
time (latent heat of crystallization of an innercore?). Existing calculations (Siegfried and Solomon, 2974; Cassen et al., I976; Solomon, I976 ,
2978a ) suggest that an iron core would solidify
during i to 3 x IO9 yr unless substantial heatsources remain in the lower mantle or core, or
perhaps if the viscosity and thermal conductivity
were incorrectly estimated. The cosmochemical
expectation of no volatile elements and no Fe
substitution in Mg-silicates actually leads to correct qualitative changes in the viscosity and thermal
conductivity, but quantitative values have not been
estimated so far. Cassen et al. discussed the possibility of an iron-sulphide shell, while recognizingthat this is inconsistent with the simplest cosmochemical models. Perhaps the assumption that U
and Th are completely or nearly partitioned into
the silicate fraction may be incorrect, especially as
the silicates have higher melting-points than iron.
Experiments are needed to test whether U and Th
may not partition significantly (e.g. 20 %) into Ferich liquid with respect to Fe-free Mg-silicates.

43

As mentioned earlier, the simplest cosmochemical model (fig. 5) for Mercury assumed terminationof accretion in the Mercury 'zone' of the solar
nebula after most of the iron had condensed andjust after Mg began to condense as forsterite. Thisapparently implies a nebular pressure greater thanc.~o -4 bar (Grossman and Larimer, I974, Fig. 7),
but uncertainty in the thermochemical data must
be considered. For ~ x io -4 bar the final temperature might be c.~36o K, which is close to thebeginning of condensation ofanorthite. If Mercury
grew from planetesimals, and if there were significant overlap of the planetary feeding zones, it
should contain small amounts of materials condensing at a lower temperature and pressure. The
estimated bulk composition of Mercury (Table IV)
assumes 9o % condensation of Fe alloyed with Ni,
Co, and C r in cosmic proportion; Ioo% condensation of A1 as MgA1204, Ti as CaTiO3, andremaining Ca as CaMgSi206; and sufficient Mg as
Mg2SiO4 to bring the mass fraction of the coredown to o.66. This simplest model has 66% ironalloy, 20 ~o diopside, 8 % MgAI204, 5 % forsterite,and o.5 % perovskite. For IOO% condensation ofiron alloy, there would be I8% diopside, 7.5%
MgA1204, 8% forsterite and o.5% perovskite.
Many otherwariants are possible, including incorporation of some Ca as CaAlzSi20 s and some Mg
as MgSiO3 if overlap occurs of planetary feedingzones. For IOO% condensation of U and Th, and
90 % condensation of Fe, the weight fractions in
Mercury of U and Th are 34 and o.I5O ppm.
Melting experiments are needed to provide mineralogical models for the Mercurian mantle, but
the simplest chemical model cannot produce asurface resembling a terrestrial basalt or anorthosite.
Because Mercury is close to the Sun the impact
velocity of late-scattered bodies is high ( > 2o km/s),
and impacts lead to violent explosions with significant loss of debris. Consequently there may be noor very little gain of CI-type material from comets
and Jupiter-perturbed bodies: indeed, impacts may
have removed a significant amount of the surface
material as Mercury was growing.
In conclusion, the arguments of Kaula (I976)
provide a warning about the application of a simple
condensation model to Mercury, and emphasize
the importance of planetesimal scattering. It is
fascinating to see how the early ideas of Urey (1952)
and Ringwood (2966b) still can be seen underlying
the more complex ideas of recent years.

44

J . V . SMITH
VENUS
And the planet of Love is on highBeginning to faint in the light that she loves
On a bed of daffodil sky.
Alfred, Lord Tennyson, Maud, I. xxn. i

There are no detailed mineralogical data on the'Queen of Heav'n, with crescent horns', but 7-ray
spectra from Venera landers indicated granitic andbasaltic compositions, thereby implying planetarydifferentiation. Although the bulk composition of
Venus is only weakly constrained the similarity of
K/U ratio for terrestrial and Venusian surface
rocks, and the approximately equal abundance ofCO2, suggest that Venus and Earth have similarbulk compositions. However, there is considerable
tolerance in the ratios of Fe, Si, Mg, S, and O. The
mineralogy at the Venusian surface must be quite
different from on Earth because of the higher
surface temperature (c.74o+_c.Io K) and thickatmosphere (coo bars = 9 x m 6 Pa) dominated by
CO2, and indeed the crust should be dominated byhigh-grade metamorphic rocks. The present
account relies heavily on a review by Hunten et al.
(I977).
Surface properties
Radar interferometry allows penetration of the
thick clouds, but even the most detailed maps
produced up to 1976 (Campbell et al., I976;
Goldstein et al., I976) have a formal resolution ofo-2o km and a geologically meaningful resolutionof perhaps c. Ioo km. Variations in reflectivity may
result from surface roughness. Deviations fromsphericity are mostly less than 6 km, perhaps
supporting Mueller's (i969) speculation (qualified
by some chemical caveats) that the higher temperature of the Venusian surface results in lower
viscosity of Venusian than terrestrial rocks. Somenear-circular features ranging from IOOto IOOOkm
diameter might be tentatively identified as craters,some ten times shallower than lunar craters. At
least two bright areas of elevated terrain might bevolcanic shields. Saunders and Malin (I977) reported that the radar-bright feature 'Beta' is 7oo km
diameter with elevation up to IO km above thesurrounding terrain. A 6o x 9~ km depression at thesummit is part of the evidence that 'Beta' is a huge
volcanic construct analogous to ones on Mars.
Linear features may be canyons or rifts. Venus isbrighter at radar wavelengths than Mercury or theMoon, perhaps indicating a thinner regolith overlying solid rock. This is not surprising since thethick atmosphere 'blankets' the surface from small
meteoroids. There is insufficient evidence to test
whether plate tectonic processes operate on Venus.

The Venera 9 camera revealed boulders litteredon a fine-grained surface. The boulders appear tobe angular to discoidal with horizontal dimension
up to 7o cm and depth up to 2o cm. At least two arelayered, and one looks like a hamburger. The
Venera IO photograph shows a curious mottled
pattern apparently of low relief suggesting finegrained material partly covering rock outcrops.
There is no obvious morphological clue to thedegradation processes necessary to 'explain thephotographs (Florensky et al., I977). Observed
wind velocities of 0.4-1.3 m/sec should be sufficient
to erode loose material because of the denseness of
the atmosphere, and chemical and aeolian erosion
may act in conjunction.
Gamma-ray measurements yielded the following
analyses:
K
Th
U

Venera 8
4.0 (I.2) w t %6.5 (o.2) ppm
2.2 (o.7) ppm

Venera 90.47 (8)3.65 (42)o.6o (I6)

Venera ~o0.30 (16)o.7o (34)o.46 (26)(Surkov et al., I974; Florensky et al., I977). Within
the experimental errors these analyses lie in theranges for terrestrial granite (Venera 8) and basalt(Veneras 9 and Io). The Venera IO rock densitometer gave a density of 2.8+o.I g/cm 3.

Atmosphere
Thick clouds of spherical droplets, I #m across,extend from 7o to 5o km above the surface. The
refractive index and infra-red absorption spectrum
are consistent with H2SO4, but other acids cannot
be ruled out as minor constituents, and might be
expected as the ultimate products of volcanic
exhalations. The yellow colour of the clouds probably results from native sulphur.
The atmospheric pressure at the surface is 9~ bar.
Carbon dioxide accounts for over 90 ~o, and thebalance (o-7 ~o?)can be attributed largely to Ar andN2, whose concentrations are still unmeasured but
have been estimated by Turekian and Clark (i 975).
Water amounts to c.oA ~ below the clouds, butsome hundred times lower above the clouds. Minor
observed constituents are CO 50 ppm, He IO ppm,
H2SO4 c.io ppm, HC1 o.4 ppm, HF o.or ppm, andO 2 ~ I ppm.
Complex chemical reactions, including catalytic
and photo-dissociation processes, are invoked toM I N E R A L O G Y OF THE PLANETS
explain the composition of the atmosphere. The
cloud tops are near 25 ~ K, and the Venus atmosphere acts as a greenhouse with respect toabsorbed solar radiation, in spite of the higher
albedo of the Venus clouds than of planetary
regolith. Radioactive heating from the Venusianinterior has only a trivial effect on the heat balanceof the atmosphere. Current atmospheric models do
not utilize the concept of continuous emission of
volcanic gases, though this must have been important for initial development of the atmosphere.

Speculations on bulk composition
Before considering the interaction of the atmosphere with the surface minerals it is necessary toreview speculations on the bulk composition of
Venus.
The simplest model has identical chemicalcompositions of Venus and Earth, except for
presumed major loss of H from Venus. However,Lewis (I972a) and particularly Ringwood andAnderson (1977) argued that the uncompressed
density of Venus is too low with respect to that ofEarth. [Note that the conversion factor from thereal density depends somewhat on the assumed
chemical composition, and the depth distributionof density, pressure, and temperature.] In particular, Ringwood and Anderson estimated that theuncompressed density of Venus is 1.9 ~o less thanthat of Earth, and that uncertainties in the pretes
have little effect, especially as Earth and Venus
apparently have similar K/U ratios. Revision of theVenus density from 5.24 g/cm a to 5.27 (Hunten etal., 1977) reduces the above difference to 1.2 ~o.
Lewis (I 972a) suggested that Venus accreted at ahigher temperature than Earth, resulting in a
mantle of Fe 2 +-free magnesium silicates, a core of
Fe,Ni metal, and a silica-rich crust similar to the
Earth's crust. Sulphur was supposed to be 'virtuallyabsent', and Lewis (1974) appealed to cometaryimpacts for supply of S to the atmosphere. Ringwood and Anderson (1977) strongly criticized thismodel with four propositions: its density is e.1 ~otoo high; the K / U ratio of surface rocks is like thatof Earth.rocks; sulphur is a major component of
the clouds, and should also be abundant in surface
rocks; and, finally, the high CO2 content of theatmosphere is not explainable by release from a
reduced mantle. Lewis has not had time to reply,
but the cumulative effect of these criticisms is rather
compelling, and difficult to refute quantitatively
though perhaps capable of being met by qualitativearguments based on late addition of CI materialand selective loss of atmospheric components.
Ringwood and Anderson preferred the Ringwood model in which Venus has the same relative

45

abundances of Fe, Si, Mg, A1, and Ca as Earth, but
in which Venus is more oxidized than Earth. The
greater proportion of oxygen lowers the density,and transfers Fe from the metallic core to FeO in
the mantle. The core (I 6 wt ~ S) amounts to 23 ~o ofthe Venus mass, and the mantle contains I3. 4 wt ~oFe giving mol Mg/(Mg + Fe) o.76. This assumptionof I6 w t ~ S in the Venusian core is inconsistent
with Ringwood's (t977) assumption that the lightelement in the Earth's core is oxygen and notsulphur, unless nebular condensates have an unexpected variation with a planet's distance fromthe Sun.
Ringwood and Anderson concluded that Venus
accreted at lower mean temperature than Earthbecause of its smaller size and gravitational accretion energy. They dismissed silicate-metal segregation in the accretion process as implausible because
it is based on undemonstrated ad hoc processes.
Nevertheless I think that metal-silicate segregationmust have taken place during the planetesimalstage, as evidenced by many properties of meteorites, and I would not rule out a lower content of Fe
metal in Venus than in Earth. I find the similarity of
K/U/Th ratios on Earth and Venus compelling
evidence that Venus and Earth accreted material
with similar volatile content. This allows a simple
explanation of the similar content of CO2 on thetwo planets (next section) and the substantial S
content of the Venusian clouds. In the next section
it will be assumed that Venusian surface rocks have
similar chemical composition to those on Earthexcept for differences caused by reaction with theatmosphere and by biological processes. If the
Ringwood-Anderson model is correct the Venusian silicate-bearing rocks should have higher
Fe/Mg than on Earth, but other models withsmaller Fe/Mg ratios are possibie.

Speculations On surface mineralogy and atmosphericcompositions
The following account largely follows the ideasof R. F. Mueller, who established the key featuresof the interaction between the atmosphere andsurface minerals. Only minor modifications wereneeded as new data became available.
Urey (1952) suggested that the high content of
CO2 in the Venusian atmosphere resulted fromfailure of silicates (e.g. enstatite and wollastonite) toreact with atmospheric COz to produce quartz andcarbonate (magnesite or calcite), perhaps because
water was not available as a catalyst. Mueller (1963)
and Mueller and Kridelbaugh (1973) considered
the rate of reactions, and in particular the latter
authors concluded that equilibration should takeplace in geologically short periods even for dry

46

J. V. S M I T H

rocks. Mueller's conclusion that all the C O 2 ofVenus is in the atmosphere, and that there is nobuffering by surface minerals was reaffirmed by
Orville (1975). Since basaltic rocks should dominate the surface, Orville re-examined possible reactions involving C O / f o r bulk compositions in theforsterite-enstatite-diopside ternary, and foundthat such compositions would occur as silicates at
741 K and pressures up to several hundred barof CO z. Acid igneous rocks (e.g. diorites andgranites) would not react with CO2 at the Venusiansurface. Several reactions listed by Lewis (I97o)
involve minerals that could not be stable at thesurface (e.g. jadeite and gtkermanite).
I now suggest that some CO2 might be stored
inside Venus. The suggested range (fig. 2) of pretes
for Venus not only intersects the field of stability ofmeionite, a variety of scapolite, but allows theassemblage orthopyroxene+Ca, Mg-carbonate to
be stable with respect to forsterite + clinopyroxene+COE, and allows calcium carbonate-silicate
minerals to be stable, as typified by the assemblage
spurrite+rankinite with respect to larnite+CO2.
Furthermore, carbon might be stored inside Venus
as graphite or diamond in the mantle, or perhaps asa solid solution with iron in the core.
Evaluation of these possibilities will require very
detailed speculation on the availability of suitablerock compositions. Especially difficult will be
evaluation of the pressure-temperature-time history of Venus and the extent of geochemical cycling
between atmosphere, crust, and mantle. At thisstage of thinking about Venus it is simplest to
assume that essentially all the CO2 is in theatmosphere where it arrived by volcanic action,
which stripped the mantle clean. Since the amountof CO 2 in the Venusian atmosphere is similar tothat in the outer zones of the Earth, it would be
surprising from cosmochemical arguments if large
amounts of CO 2 were hidden below the surface ofVenus. If the surface temperature of Venus were todrop below c.68o K, Mg-silicates should begin toreact with CO z with the ultimate possibility of athin atmosphere with no greenhouse heating.
The height distribution of CO in the Venusian
upper atmosphere was modelled by photochemical
reactions and vertical mixing (Sze and McElroy,
I975). If the CO/CO2 ratio of the lower atmosphere
were shown to depend on reaction with surface
minerals, or on volcanic action, there might beimportant implications for the oxidation state
either at the surface or inside Venus.
The partial pressure ofc.o. I bar H 2 0 in the lower
Venus atmosphere might be explained either by
input from continuing volcanism and cometaryand meteoroid debris, or by buffering from surface
minerals, or by some hybrid non-equilibrium state.

Although there is a psychological tendency to thinkthat HzO has left the Venus interior along with theCO2 and that it must have continued to moveoutwards rather than returning, possible storage in
the interior should be considered. The present
content of H 2 0 in Venus is hard to estimate though
an upper limit can be calculated fairly accurately.
Even more uncertain is the primordial content
because of the difficulty of quantifying the potential loss through the atmosphere and ionosphere.
All known minerals containing H 2 0 molecules
are unstable on Venus, and many hydroxylated
ones are unstable. Of the major hydroxylated
minerals, only amphibole and mica could be important stores (Mueller, I964). For peridotitic rocks,
amphibole (of appropriate composition) could bestable to I5-2o kb for low Fe/Mg (cf. figs. 2 and 8),
but to lower pressures as Fe/Mg increases.
Phlogopite-rich mica could be stable to a higher
pressure whose limit is difficult to ascertain becauseof the near-asymptotic approach of a prete to thecurve for breakdown (fig. 8). On Earth, amphibole
may occur down to c.Ioo km and phlogopite toc.I5O km, and probably these depths are notexceeded on Venus. Of course, even mica andamphibole may not occur on Venus because volcanic action may have driven the water into theatmosphere, where it was lost through the ionosphere before it could be reacted at the surface andsubducted downwards.
It turns out that mica can be a significant store
even if Venus has no more K than Earth, as isrequired by cosmochemical models. Paragonite
and margarite cannot be significant contributors,
and all micas must be K-rich with K/HeO c.2. If
Venus contains I3o ppm K, mica could store c.6oppm H20, which is about one-tenth the amount of
H 2 0 estimated for Earth (Smith, I977). Of course,much or all of the K on Venus might be tied up infeldspar, thereby reducing the above estimate. The
maximum amount of amphibole can be estimatedby placing a stability limit of c.6o kin, corresponding to c.4oo ppm H : O in Venus. Thus the outside
limit of c.5oo ppm H 2 0 is similar to the estimatedc.6oo ppm for Earth, but many arguments can beinvented to suggest that Venus has much less than
5o0 ppm H20.
Loss of water from Venus depends on volcanic
emission and transport upwards in the atmosphere.
Models for generation and loss of water are
discussed in detail by Mueller (I964, I97O), Fricker
and Reynolds (I968), Palm (1969), Rasool and
DeBergh (I97O), Turekian and Clark (I975), Prinn(1975), Walker (1975), and Sill (1976). Loss of waterfrom Venus cannot be estimated a priori because it
depends on complex diffusion-controlled processes. Uncertainty in the transport mechanism inM I N E R A L O G Y OF THE PLANETS
the atmosphere, especially in relation to a sulphuricacid barrier, is particularly severe.
Volatiles can be modelled by analogy withmeteorites and Cx material has been used. The
simplest reaction relates organic material, represented by CH2, and magnetite, Fe304: CH2+
3Fe304 ~ CO2 + H 2 0 + 9 F e O .
This requires
equal abundances of CO2 and H20 on Venus,
except for H 2 0 lost through the atmosphere. If
H 2 0 were lost in toto there is no irreversible change
in the oxidation state, but the more likely prospectthat H is lost more efficiently than O requires thatthe remaining O must be taken up by the crustal
rocks. Loss only of H corresponds to oxidation of
Fe 2+ to Fe 3+ producing some Io km of magnetite
or c. IOOkm of oxidized basalt. Sill (I976) suggested
that volcanic gases on Venus were like those on
Earth in containing roughly equal molar proportions of H20, CO2, and SO/, and he proposed thereaction SOz + H 2 0 ~ SO3 +H2~'. The S O 3 w a s
incorporated into sulphate by reactions such as
CaMgSizO6 + 2SO3 ~ CaSO4 + MgSO4 + 2SIO2.
Complex models for the atmosphere have beendeveloped with production of large amounts of O2
in the early atmosphere (Walker, I975) and use of
the 'runaway greenhouse' mechanism (Rasool and
DeBergh, I97O).
Mueller's (I965) investigation of the stability ofsulphur compounds on Venus is partly obsolete
because of revisions in the atmospheric composition, but the basic features remain intact (Sill,I976). At the surface, anhydrite and the feldspathoids sulphate-scapolite (Goldsmith, I976) andhaiiyne are potential sinks for S, and iron sulphides
are unstable if the oxidation state is controlled bymagnetite and hematite. Sill (i976) also considered
the possibility of a reduced state for the Venusiansurface in which troilite and perhaps pyrite arestable, thereby allowing for a significant partialpressure of S vapour. This S vapour might explain
the yellow colour of the clouds. However, I prefer
the concept of an oxidized surface, which could betested experimentally by magnetic measurements(cf. Mariner experiments on Mars).
The small but non-zero mixing ratios of HC1 and
HF in the atmosphere require detailed explanation(Mueller, I968; Sill, I976). Several minerals including fluorite, mica, and amphibole offer potential
sinks or buffers for F, while scapolite and cancrinite

47

could serve for C1. Perhaps continuing volcanic
activity is supplying HCt and HF to the atmosphere, and perhaps there are substantial amounts
dissolved in H2SO4 in the clouds.
Concluding remarks

Major uncertainties remain about the mineralogy, petrology, and geochemistry of Venus, and it
will be very difficult to obtain detailed mineralogical information about the surface because of
the strain placed upon instrumental design by thehigh temperature. Perhaps magnetic and spectralreflectance studies will provide useful data, especially on the oxidation state. Very useful would bedetermination of the extent of volcanic activity. If
radar studies of high spatial resolution indicate
volcanic morphologies, it might ultimately be possible to fly a mass spectrometer over possible cones
to test for emission of gases. In the meantime,mineralogists must rely mainly on further study ofthe composition and height distribution of theatmosphere, preferably by high-resolution mass
spectrometry.
The thermal evolution of Venus may be complicated by capture of one or more moons (as invoked
to explain the slow retrograde rotation: Singer,
97o; French, 1971; Ward and Reid, 1973). Because
the planet is so large and because its K/U ratio islike that of Earth it is tempting to invoke thepresence of plate tectonic processes continuing to
the present time. If so, the volcanic gases should be
providing information on the interior. Can it beargued that the paucity of water in the atmosphereresults from a dry interior, or is water removed bysurface reactions as fast as it is supplied? It is
certain that low-grade metamorphic rocks cannot
exist on Venus, and it is likely that the pretecorresponds roughly to the boundary between
medium and high grades of metamorphism as
defined for terrestrial rocks (Winkler, I974,
Fig. 7--2). Does anatexis occur? Can most crustal
rocks be loosely described as granulites? Whatever the answers, Venus must have much simpler
mineralogy and petrology than Earth andmany abundant minerals on Earth (including
muscovite and gypsum) cannot occur on Venus,
while others (including microcline) can be stable
only at shallow depths.

48

J . V . SMITH
MARS
And yonder all before us lie
Deserts of vast eternity.
Andrew Marvell, To his co), Mistress

The spectacular Mariner and Viking missions
revealed the major features of the atmosphere andsurface of the 'red' planet of war and death, but
provided only tantalizing hints of the mineralogy,
petrology, and chemistry of the sub-surface. At thetime of first writing, debate was just beginning on
the complex results from the Viking missions (V~,
Chryse Planitia; V2, Utopia Planitia), and thepresent account is based largely on articles inScience (27 Aug., I Oct., I7 Dec. 1976 ). Reviews of
the geology were given by Mutch and Saunders(I976) and Mutch et al. (I976); of the generalproperties of Mars by J. B. Pollack in Scientific
American, Sept. I975; and of the atmosphere by
Moroz (i 976). Implications for the evolution of the
Earth from Martian observations were discussed
by Rasool et al. (1977). The September issue of J.
Geophys. Res. appeared during revision, as did anarticle on the abundance of volatiles on Earth andMars (Anders and Owen, I977).

Surface morphology

Bearing in mind the difficulties of interpretationof physiography and the lack of an absolute timescale for Martian stratigraphy, the following summary provides a temporary basis for mineralogical
speculation.
The oldest units are the undivided cratered
terrain, which covers rather more than one hemisphere centred on c.55 ~ S, 325 ~ W, and the undivided plains, which occupy the remaining antipodal fraction. Almost certainly the heavy cratering of the undivided cratered terrain results, by
analogy with the lunar highlands, from the period
earlier than c . - 4 . o x io 9 yr. Lower abundance of
craters in the undivided plains can be explained by
obliteration of craters by extensive volcanic
activity. Whereas most craters of the northern
plains have sharp rims, those in the southern terrain
have a range of sharpness suggesting considerably
greater erosion for older craters. There is no
obvious explanation for the asymmetry of the earlyMars surface and its angular displacement from thepresent polar axis.
Huge near-circular structures, including the
Argyre, Isidis, and Hellas basins (the largest withdiameter 2ooo km and depth 4 km) have a size andfrequency that would allow matching with lunar
and Mercurian impact basins (Wood and Head,I976), probably providing a time marker near
- 4 • lO 9 yr.
Several types of plains have been distinguished,but it is uncertain whether they are surfaced byvolcanic flows or impact debris.
Huge masses of volcanic rocks definitely poured
out of Mars (Carr et al., I977b). The saucer-shaped
Alba Patera, 16oo km across, may be the largest
central-vent volcano in the solar system. Its very
low profile is not matched by volcanic structures on
Earth, Moon, or Mercury. The Tharsis andElysium plateaux are capped by enormous shield
volcanoes, typified by Olympus Mons whose cone(with a summit caldera 8o km across) rises 25 km
above a scalloped escarpment 6oo km across and
6 km high. Surrounding the cone is a ring of plains
lOO zoo km wide and an outer incomplete girdle ofgrooved terrain 2oo-7oo km across. Perhaps thecomplex data on volcanic landforms will be fitted
into an over-all sequence beginning with planetwide volcanism in an era of heavy impacts andisostatic adjustment, and terminating with extrusion of lavas in uncompensated shield volcanoes
that grew over stationary hot-spots in the mantle
for long periods of unknown duration (io s toio 9 yr?).
Although the circular structures are particularly
eye-catching, the major volume of volcanic rocks
probably occurs in inconspicuous plains. The
morphology of most lava flows that can be recognized as individual units indicates a flow volume
much greater than on Earth and a low viscosity.
Flows near the summit of the Martian shield
volcanoes are shorter than those at lower elevation,
perhaps indicating higher viscosity of late lavas.
Three types of lava channels (Masursky et al., I977)
were distinguished from three types of channels
attributed to flow of water.
These latter channels were classified into broadchannels abutting areas of chaotic terrain (Carr etal., ~977a, Fig. 7), possibly caused by catastrophic
melting of frozen ground by subsurface heating;sinuous channels of widespread distribution attributed to underground melt-water produced frompermafrost by climatic warming; and dendritic
channels, which terminate abruptly on plains or
crater walls, and which may result from rainfall.
Carr and Schaber (I977) attributed the following
features to extensive permafrost in the ancientcratered uplands: fretted terrain, chaotic terrain,alases, canyons, patterned ground, and unusualM I N E R A L O G Y OF THE PLANETS
morphology of impact craters (see also Carr et al.,I977c ). Although water, derived either from permafrost or rainfall, is a probable scouring agent for
many Martian channels, rifting coupled with landslides and aeolian erosion may be responsible forother linear features.
The enormous region ofValles Marineris, whichspans one-quarter of the equator, was attributedby Blasius et al. U977) to extensional tectonic
activity followed by a variety of erosional anddepositional processes. Large blocks have shiftedvertically and tilted with respect to each other.
Massive landslides, perhaps lubricated by fluid,
modified the valley shapes, and the whole setof processes probably lasted over Io8-Io 9 yr into
recent times.
Huge dust clouds periodically obscure Mars as
spring ends in the so.uthern hemisphere. The
layered terrain near 75 ~ N and S consists ofnumerous uncratered terraced layers with an
unconformity, all overlain by deposits of permanent ice (Cutts et al., I977). Its origin is notunderstood, but it may result from episodic deposition of dust at the edge of a mid-latitude storm
belt that was affected by long-term climaticchanges. Major dune fields adjacent to the layered
terrain, and minor dunes at the Viking sites,
testify to aeolian processes.
Thermal and albedo mapping revealed two
major types of Martian surface material and acovering of fine particles on bright surfaces (Kieffer
et al., I977b). Frey (I974) suggested that there were
two distinct types of crustal material on Mars bysimple analogy with Earth and Moon, but there areno specific chemical indications. The causes ofalbedo changes on Mars are not understood, anddust deposition may be only one factor.
Three types of aerosols were detected over theViking landers (Pollack et al., I977): ground fog
during late summer nights, equivalent to 2 #m of
precipitable H 2 0 ; high-level ice clouds in theautumn; and mineral particles (o.4 pro, non-spherical, rough), whose imaginary refractive index fits
well with t o + 5 ~ o of magnetite mixed with anunidentifiable material. Pollack et al. suggested
that a coating of CO2 would result in more rapid
fall-out of mineral particles and associated HzOice over the winter poles.
The permanent portions of the polar caps consistof water ice and dirt (Kieffer et al., I977a ), and areaugmented by small seasonal ice caps containing
H 2 0 and CO2 ices, which migrate between thepoles by vapourization, atmospheric transport, andcondensation. The thickness of the ice caps is
unknown. Dzurism and Blasius (x975) discussed
the effects of higher surface elevation and lower
pressure of the southern cap, and suggested that the49

caps are I-2 km thick in the south and 4-6 km in
the north (see also Farmer et al., I977 ).
Dating of surfaces from crater counts is highly
controversial because of lack of calibration.
Whereas Neukum and Wise (I976) argued forcompletion o f the major volcanic episodes in thefirst I. 5 • Io 9 yr with Olympus Mons completed by- - 2 . 5 X I O 9 yr, Hartmann (I977) suggested a much
longer volcanic history with Olympus Mons only
o.o6-o.6x ~o9 yr old, and the channel systemformed a t - 3 t o - o . 3 • IO9 yr. Whatever the absolute time scale the growth period of the volcanic
shields on Mars is much longer than on Earth(Blasius, I976 ). See also Chapman and Jones(I977).
Small craters have sharp rims suggesting low
erosion in recent times (IO8-Io 9 yr?) in spite of thedust storms. The great thickness of layered deposits
near the poles and the high rate of obliteration ofearly craters were attributed to an early episode ofstrong erosion, perhaps associated with a thickatmosphere (Jones, I974).
Topographically, the northern plains are lower
by about 2 km than the southern cratered terrain.
Mars is mostly at isostatic equilibrium, except for
the Tharsis plateau whose positive anomaly may berelated to the negative anomaly of the adjacent
Chryse region by underlying mantle convection.
There is no evidence of plate-tectonic processes, but
major fractures probably mark incipient rifting.
Panoramic views at the Chryse and Utopia
landing sites (Mutch et al., ~977a, b) revealed
boulder-strewn deserts. Greeley et al, (I977) interpreted the geology of Chryse Planitia in terms ofa basement of heavily cratered ancient terraincovered first by deposits of aeolian, fluvial, andvolcanic types, and more recently by d e p o s i t s ,probably of volcanic rocks and impact debris.
Binder et al. (I977) attributed the view from the VI
lander to volcanic terrain exposed as bed-rock onsome crests, upon which reside numerous blocks ofseveral rock types, whose pitting indicates a grain
size in the mm to cm range, and fine-grained
sediment and duricrust. The view from the V2
lander is also dominated by blocks lying on a finegrained surface. Mutch et al. (1977c) suggested that
Utopia Planitia has volcanic domes, and that its
sediments and dunes are part of the mantle extending northward from 35 ~ latitude. The blocks have
similar surfaces whose pits of mm-cm size were
thought to represent vesicles rather than theweathering of soft crystals. A polygonal network
was attributed to ice wedging of contraction cracks.
Pebbles sampled by the X-ray fluorescence experiment have similar composition to soil except for
extra S, and probably represent fragments ofduricrust t-2 cm thick. No chemical data were

J. V. SMITH

50

obtained for the boulders and the surface homogeneity and even distribution argue against attribution to the ejecta blanket of nearby Mies crater,
Fragmentation of lava flows followed by sedimentary and aeolian processes may explain theboulders.
Early attempts to understand the surface ofMars relied on its red colour, but Huck et al. (I977)
discovered that the Martian surface is actually
yellowish brown with local change to grey. SomeV I rocks appear dark grey, but most soils and rocks
range from dark-brown to strong yellowishbrown and even moderate olive-brown. The reflectance has a strong band centred at c.o.93/~m, which
Huck et al. prefer to match with nontronite rather
than limonite. Binder et al. (1977) prefer to retain
the suggestion by Binder and Cruickshank thatlimonite is responsible, whereas Hargraves et al.
(I977b) prefer maghemite (~-F%O3). Whether theyellow-brown colour merely results from a surface
patina on grey rocks (basaltic?) and clay minerals is
not certain, and neither is it clear whether air-borne
dust and aeolian sediments differ in colour.
Atmosphereand volatiles

The pressure at the surface ranges from about 5
to IO mbar, decreasing with elevation and changing
slightly with season and time of day. Mixing ratiosfrom various techniques are CO 2 c.95 ~o, N2 2.7 ~o,
Ar 1.6~o, 02 o.13 ~o, CO o.o7 ~ , H 2 0 variable but
c.o.o3 ~o, Ne c.2. 5 ppm, Kr 0. 3 ppm, Xe o.08 ppm,
and 0 3 o.o3 ppm (Owen et al., 1977). The watercontent of the atmosphere varies greatly, but,
although sometimes high enough to produce local
mists, amounts only to t km a for the wholeplanet.
Enrichment of 15N over 14N in the Martian
atmosphere at c.13o km height (15N/14N o.oo64)
over Earth (o.oo37) can be explained by differential
loss of the lighter isotope from the upper atmosphere thereby suggesting a substantially higher
concentration of nitrogen in earlier times (McElroy
et al., 1977b ). Ratios for l s o / 1 6 0 (o.oo2o) and13C/12C (o.ot I5) match those for the Earth within
experimental error, and imply large accessible
reservoirs (at least equivalent to several bars of
atmospheric pressure) of O and C for exchangewith the atmosphere if the differential escapemechanism used for N 2 is not to produce detectablechanges for O and C (Nier et al., 1977).
Mixing ratios and isotopic ratios of the raregases place important restraints on models foroutgassing and atmospheric escape. Owen and
Biemann (1977) found that 4~
is 2750 ___500,
which is ten times greater than on Earth. Owen etal. (I977) detected Xe and Kr, and found that lZ9Xe

is much more abundant than 132Xe and t31Xe.
They found that non-radiogenic noble gases in theMartian atmosphere have the same relative abundance as in the terrestrial atmosphere, but with IOOfold lower abundance when scaled to planetary
mass. The deficiency of Xe in the Earth's atmosphere with respect to chrondritic abundance has
been attributed to storage in shales and othersediments, and the probable deficiency in theMartian atmosphere may similarly be attributed tosedimentary reservoirs.
Some suggestions published before the Viking
mission are now obsolete because of the erroneous
assumption of twenty-fold higher abundance of Ar
in the Martian atmosphere, but their general ideas
are still useful (e.g. Fanale, 1976; Levine, I976;
Owen, I976 ). Current speculations involve fivefeatures: a denser early atmosphere, perhaps denseenough (o.I to I bar?) to allow heavy rainstorms;
considerable exospheric loss of light gases to explain the 15N/14N ratio; reaction of C- and Obearing species with crustal reservoirs; migrationof much of the early atmosphere into the regolith,either by freezing (permafrost), adsorption, or
chemical reaction; and only partial outgassing,
especially of 4~ produced from the decay of 4~
The climate and atmosphere may have changed
considerably because of either change of the energy
output of the Sun (Hartmann, I974) or theobliquity and eccentricity of the Martian orbit(Ward, 1974), and these changes even may havebeen complicated by greenhouse-type mechanisms(Sagan et al., 1973).
The photochemical reactions in the atmospherehave profound implications for the surface mineralogy. The following section is based on a lecture by
D. M. Hunten in November 1977 in which hefollowed up ideas in Hunten (1974), McElroy andKong (1976), Kong and McElroy (I977), andMcElroy et al. (I977a).
The key is photolysis of H 2 0 in the atmosphereby solar ultraviolet radiation producing H andOH radicals. Various photochemical reactions,whose details are irrelevant here, then produce H202
and other species including HO2. The hydrogen
peroxide freezes out under almost the same conditions as H20, while HO2 condenses at a lower
temperature. The estimated production of approximately lO 17 molecules at H202 per cm 2 per year
would suffice to explain the absence of organic
molecules in the soil and the evolution of gases inthe biologically oriented experiments on the Viking
landers (Biemann et al., 1977; Horowitz et al., 1977;
Oyama and Berdahl, 1977; Ponnamperuma et al.,1977)- Ironically this photochemical scheme resultsin water destroying the pdssibility of life on Mars
rather than fostering life as on Earth!

M I N E R A L O G Y OF THE P L A N E T S
Mineralogy and petrology

Weight restrictions limited Viking instrumentation to an energy-dispersive X-ray fluorescence
spectrometer that could not detect elements lower
than Mg. Furthermore, the poor energy resolutionof the proportional counter (resulting from theimpracticability of using a high-resolution crystal
cooled to liquid N2 temperature) caused strong
overlap of contributions from elements adjacent in
the Periodic Table. The early analyses of fines fromboth Viking sites (Clark et al., I976 ) were remarkably similar in having high Fe (I 3 --+2 wt ~o), moderate Mg (5 -+ 2-5), Ca (4 + o.8), and S (3 -+ o.5), low A1(3+o.9) and K ( < o.25), and very low trace elements (e.g. Sr c.8o ppm; Y c.6o ppm; Rb < 3o ppm).
Other elements are: Si 2o.9+2.5; C1 o.7+o.3; Ti
0.5__+o.2. Computer modelling led to a chemical
mixture of 81 ~o Fe-rich clay (59 ~o nontronite; 22montmorillonite), 12 ~ Mg-sulphate (kieserite), 6
calcite, I ~ halite, and o.5 ~ rutile. Some I to 7 ~o ofthe fines became attached to magnets (Hargraves
et al., I977b), thereby indicating some type ofmagnetic iron oxide. Although details are uncertain
the analyses were attributed to the weathering
products ofmafic igneous rocks (Baird et al., 1977).
Unpublished analyses of 2-12 mm fragments at V2
yielded 2 5 - 5 o ~ more S than for fines, but nosympathetic increase of Ca, A1, and Fe. A sample offines, protected from aeolian modification by burial
under a rock, gave lO-15 ~ less Fe than the earlierfines. The correction procedure for the analyses
assumes that the material is spatially uniform. A
thick skin of iron oxide coating silicate material
would yield falsely high Fe and over-all sum of
calculated oxides, and Clark et al. (I977) placed an
upper limit of o.24 #m on such a skin. Baird et al.
(1977) concluded that the Martian surface has anabundance of mafic parents for subsequent clay
formation, a lack of large exposed alkali-richgranitic materials, and a sufficiency of water and ice
to form clays. Schonfeld (I977, oral presentation,
Lunar Science Conference) argued that theanalyses can be well represented by a mixture of Ci
meteorite and basalt, but details have not appearedin print.
Using a gas-chromatograph-mass-spectrometer, Biemann et al. (1977) did not detect organic
molecules at the IO-3 ppm level, and observed
release only of H 2 0 and CO2 upon heating fines to
5o~ 2oo ~ 35 o~ and 5oo~ for 3~ seconds. From o.i
to I wt ~ H 2 0 and 5o-7oo ppm CO2 were released,
but identification of the mineral hosts is unclear
because of uncertainty in the kinetics. Perhaps thewater was released from hydroxylated silicate or
hydrated sulphates or both, while the CO2 might
have derived from carbonates less stable than51

calcite (e.g. magnesite or siderite). Absence of Sbearing gases rules out presence of weakly bonded
sulphides, and the most likely hosts for the sulphur
are strongly bonded sulphates not containing substantial A1, Ca, or Fe.
Revision of the colour of the Martian surface
from reddish to yellow-brown changed the mineralogical implications. Instead of favouring a thinfilm ( < 0.25 pm) of hematite on surfaces, the colour
favours maghemite (7-Fe203). Hargraves et al.
(I977a) had preferred to interpret the magnetic
material at the Viking sites as magnetite (possibly
titanian to explain the o.5___o.2~ Ti in the XRF
analyses), but later (I977b) switched their preference to maghemite.
Toulmin et al. (I977) pointed out that nontronite
breaks down to maghemite and siliceous material
under oxidizing or inert atmospheres, and thatmaghemite does not transform to hematite below
8oo ~ Huguenin (i976 and earlier references)
showed experimentally that ultraviolet radiation
produced photo-stimulated oxidation at the surface of magnetite and Fe-bearing silicates under
dry conditions in the presence of O2 gas. In thepreceding section the photochemical conversion of
HzO into H202 was described, and reaction withhydrogen peroxide (and perhaps other 'superoxides') might be more important than direct
photo-stimulated oxidation at a mineral surface.
Furthermore, burial and later exhumation of sediments might result in complex low-grade reactions
dominated by surface reactions, perhaps of acatalytic nature.
Spectral data for dust clouds need re-examination. Aronson and Emslie (1975) interpreted infrared data for Martian dust in terms of a strongly
polymerized silicate, preferably feldspar and possibly a layered mineral, and ruled Out goethite fromabsence of 8oo and 9oo c m - 1 bands. Perhaps thesuggestion by Pollack et al. (1977) that magnetite is
involved may need revision: maghemite might also
be a suitable candidate because of its similar
crystal-chemical properties.
The S content of XRF analyses of Viking fines
can be interpreted in terms of hydrated Mgsulphate(s) and the chlorine content (o.7 + 0.3) may
result from salt(s). Malin (I974) suggested that salt
weathering, which is believed to occur in Antarctic
dry valleys, may occur on Mars. The salts may
derive from volcanic gases, and the pressure andtemperature, especially for equatorial latitudes, areconsistent with temporary formation of brine(Ingersoll, 197o ). Malin also pointed out that somequartz diorites from Antarctica had weathered tobrown red.
Whether melting of the clathrate CO2.6H20 is
responsible for floods is disputed (Milton, I974;

52

J. V. SMITH

Peale et al., I975) but the clathrate was predicted tooccur in association with water ice in polar regionswith temperature below c.i53 K (Miller andSmythe, 197o; Dobrovolskis and Ingersoll, 1975).
The conditions which govern the melting or
evaporation of water on Mars are complex andapparently strongly dependent on a dust cover(Farmer, I976).
Temporarily, the simplest synthesis of the clues
to the mineralogy of Mars is that the dominant
extrusive rocks were Fe-rich 'basalts', thereby explaining the low A1 in the XRF analyses, and theapparently low viscosity indicated by flow morphology; that the weathering of the 'basalts' occurredin a hydrous, oxidizing environment producing Ferich clay minerals (perhaps mainly nontronite)coated with maghemite and perhaps other ironoxides or hydroxides; that S- and C-bearing gases
were incorporated into sulphates and carbonates(probably mainly of Fe and Mg); and that various
species, especially H20 and CO2, did not react toform minerals, but formed ices, which temporarily
melted into brines or even evaporated. Perhaps
most of the material analysed by XRF ultimately
derived from dust clouds of weathered volcanic
rocks.

Geophysical and geochemical models
Unfortunately the seismic experiment on theViking missions has not recorded any event strong
enough to provide constraints on the Martianinterior, and geophysical modelling is based merelyon the density and the moment-of-inertia. Revisionof the latter from C/MR 2 =o.377+_o.ool too.365 + o.ooI (Reasenberg, 1977) resulted in majortransfer of mass towards the planetary centre, but
the history of the lunar moment-of-inertia does notinspire confidence in the latest Martian value.
Cosmochemical models for the planets are also
highly uncertain, and present data on the volatiles
on Mars are perhaps more easily interpreted interms of a volatile-poor than a volatile-rich planet(Anders and Owen, I977). Nevertheless the earlymodels for Mars are important for construction of
better models.
Simple models for condensation of the solar
nebula lead to proposals that Mars should have afull complement of'refractory' and 'silicate' groupsof elements, a higher content of S and oxidized ironthan Earth, and a greater content of volatileelements. Specifically, Lewis (I972a) proposed that
Mars has a core dominated by FeS, perhaps with alittle Fe metal, and a mantle with Fe/Mg c.1. Healso proposed that Mars accreted well within thecondensation field of tremolite, and estimated one
H20 for every 2Ca (i.e. H 2 0 c.o.3 Wt~o of bulkMars). Anders and Owen (I977) concluded that
Mars has only 3 ~ of the Earth's complement ofvolatiles (proportioned to planetary mass), andobtained them from the same source. This apparently rules out the Lewis model for Mars. Perhaps
the volatiles are derived from late impacting
planetesimals that condensed in the outer part ofthe solar system, and perhaps the lower escape
volocity of Mars resulted in retention of only a few
per cent of such material. If so, Mars may haveaccreted mostly from bodies formed from the solar
nebula at temperatures higher than that for condensation of tremolite. Perhaps the non-volatilecomposition of Mars is quite similar to that ofEarth as befits the similar values for uncompressed
density, but perhaps the degree of oxidation andsulphidization is higher by analogy with the simple
Lewis model. It is also possible to modify the ratioof refractory condensate to Mg-silicate and metal
fractions between the two planets, as in the Ganapathy-Anders model for the Earth. With thesefactors in mind the existing models can be reviewed.
Using the old moment-of-inertia, Anderson(1972) modelled Mars with several mixtures of
meteorites thereby ending up with substantial S in
the core. Varying the core composition from pure
Fe to pure FeS resulted in the core radius changing
from 122o to 2o4o kin. The zero-pressure density ofthe mantle was limited to 3.54-3.49 g/cm3, implying
24-2I Wt~o FeO unless some Fe metal remains in
the mantle. For all chondritic models Mars has asmaller and less dense core than the Earth and adenser mantle. The bulk Fe content is 25-8 wt ~.
Binder and Davis (1973), following Binder(1969), used new physical data for Mars, together
with the possibility of S in the core, to place revised
limits on the core and mantle. Particularly important is the use of an isostatically compensatedequatorial bulge of 8 km to yield a crust (assumed
2.8 g/cm 3) of thickness ranging from 24 to 65 km
depending on the mantle density and the crustal
thickness at the poles. Many complex possibilities
are allowed, but for a core density near 7.1 g/cm 3(corresponding to about 3o ~ FeS in the core), thecore is about 1250 km in radius and contains c. lO ~oof the mass while the mantle has olivine composition (Fo65_ 8o)- If the crust is not fully compensated isostatically there are greater tolerances.
Earlier, Ringwood and Clark (i 97 I) investigatedthe consequences of Ringwood's model for derivation of Mars from C i-meteoritic material, but moreoxidized than Earth and ordinary chondrites.
Using a simplified CI composition that ignored allvolatiles and S, they considered three mantle-coremodels: (i) all Fe in magnetite and all silicates Fefree; (ii) half Fe as FeO incorporated in olivine
and pyroxene, and half as magnetite; and (iii) most

MINERALOGY OF THE PLANETS
Fe as FeO and little as Fe20 3. All models could bemade to yield the observed density, but all had too
high a moment-of-inertia even for the old value. In
order to decrease the moment-of-inertia a crust ofdensity 3.o g/cm 3 was added. Finally the probablepresence of FeS was considered. Ultimately thefollowing complex model liB was proposed: deptho-c.5o km crust; 5o-11oo upper mantle of olivine
and pyroxene; I IOO-I65O lower mantle of spineland garnet phases of olivine and pyroxene composition; 165o-I85O outer core of molden FeS-O; I85O-3388 inner core of Ca-ferrite, a highpressure form of magnetite composition.
Johnston et al. (I974) set the internal structure ofMars in the context of the possible thermal state.
Geophysical parameters were chosen in order toobtain an early differentiation producing a core,mantle, and crust, a prolonged partial melting of
the mantle producing extensive volcanism, and asurface rigidification to zoo km depth in order tosupport the present non-hydrostatic shape. They
did not rule out present melting of the core, as had
Binder and Davis, because of new evidence for aweak dipole magnetic field (but see Dolginov, 1978,
for controversy with C. T. Russell). In order toobtain early melting of Mars (before - 4 x I o9 yr), it
was necessary to assume, even for the low-melting
Fe-FeS eutectic, either a high accretion temperature, or a high content of radioactive elements, orboth. Johnson et al. assumed o.o3 ~ ppm U, which isover twice the value expected for the Lewis modelof Mars. For rapid accretion (lO 5 yr) of material at
273 K, core formation occupied the period c. - 4 . 2
t o - 3 . I X 1 0 9 yr, and melting of dry silicate begana t - I . 3 x IO9 yr at 6oo km depth. For material
accreted at 773 K, core formation began almost
immediately and was completed a t - 3 . 6 x l O 9 yr
while dry silicate began to melt at - 2.6 x io 9 yr. As
discussed earlier, there is a general problem offinding heat sources for the early melting of planets,
and there are severe problems in attempting tomodel early core formation on Mars if the accretedmaterial is low in U (c.o.ol 3 ppm) and if it accreted
slowly at a low temperature. It seems very difficult
or perhaps impossible to explain the surface
morphology of Mars and its present geophysicalproperties without invoking early differentiation,and Solomon (I978a) argues that the surface
morphology of Mars requires an originally cold
interior coupled with restriction of early differentiation to the outer part. If this is accepted theanalysis by Johnston et al. must be a good guide to
the thermal development of Mars, except that there
may be a different distribution of heat sources atthe beginning, and a smaller production of radioactive heat. Whatever the details, one can imagine
an early differentiation involving separation of a53

heavy Fe-FeS eutectic or near-eutectic liquid and alight hydrous basaltic liquid, and a later differentiation involving partial melting at higher temperature of a mantle of dry peridotite.
Johnston et al. modelled Mars with a 50 km crustof density 3.o g/cm 3, a mantle of pyrolite withadded FeO to bring the density up to 3.74 g/cm3,
and a core with 85 wt ~ Fe and 15 ~o S. This results
in a mantle with 73 ~ olivine (65 mol ~o Fo), 25 ~ogarnet (grossular 53 almandinel 9 pyrop%8; wt ~o)and 2 ~o Fe,Mg-oxide. The mantle composition is
SiO2 34.6 AlzO3 2.7 Fe203 0.3 FeO 29. 5 MgO 28.7
CaO 2.4 K 2 0 o.i wt ~o- Use of a revised momentof-inertia by Johnston and Tokstz (1977) resulted
in movement of Fe from the olivine (now 75 molFo) into the core (15oo-2ooo km radius; more Fetich than the Fe-FeS eutectic) giving an estimatedbulk composition close to that in Table IV.
It is obvious that there is considerable tolerance
in models for Mars, and that the key factors are theassumed concentrations of S and O in the core, a n d
the degree of differentiation of Fe between the
mantle and core. Uncertainties in the pressuretemperature distribution are not as important.
Because of the high S content of the fines at theViking site, I prefer to assume that the volcanic
gases and volcanic rocks derived from sources
containing substantial S, and hence I prefer modelsof Mars in which S is a major element, perhaps
largely concentrated in a core but with someremaining in a silicate-rich mantle and crust.
Recent preprints by Okal and Anderson (1978),
McGetchin and Smyth (1978), and Solomon(I978a, b) extend the above discussion.
Siever (I974) compared the differentiation ofMars and Earth with particular emphasis on
geochemical transport. He pointed out thatweathering does not occur simply because amineral is brought to low temperature, but requires
reaction with water and dissolved acids. Reactions
in feldspar are governed by diffusion through analteration rim, and it is difficult to estimate a p r i o r i
the rate of alteration on Mars. He suggested that
Huguenin's (I 976) mechanism of photo-stimulatedoxidation of magnetite might also be ultimately
controlled by diffusion. Huguenin suggested that
surface scaling would greatly enhance the rate ofweathering. Booth and Kieffer 0978) concluded
that carbonate minerals can form even under arid
conditions. There are many complex problems, andit seems impossible to make detailed predictions on
the mechanisms and extent of weathering on Mars,
especially in view of the relation to photochemical
processes in the atmosphere. Siever emphasizedthat there is no evidence for vertical motions
on Mars large enough to allow recycling of sediments followed by volcanic activity. Unlike Earth

54

J. V. SMITH

sediments, all the Martian sediments should be stillat the surface or covered only by other sediments orlava flows. Volcanic emissions from Mars should
enter the atmosphere, react with surface minerals,
and remain mostly trapped there forever. Wind
transport should redistribute the altered minerals
from one place to another, and volcanic andclimatic heating should cause local melting of frost,and perhaps some devolatilization of weakly
bonded minerals, but there should be no major
geochemical cycles as on Earth.
For the atmosphere Siever mentioned thermal
escape of hydrogen and helium, non-thermal
escape of oxygen (Liu and Donahue, I976), andabsence of S- and Cl-species. He suggested thatmost S occurs on Mars as solid sulphide, andpointed to the stability of pyrite in terrestrial
Precambrian sandstones before bacterial oxidation
became important. However, the dissociation of
H 2 0 in the Martian atmosphere (McElroy andKong, i976 ) and at mineral surfaces (Huguenin,
I976) should lead to more effective inorganic
oxidation of sulphide than on Earth, and indeed theViking experiments are more easily interpreted interms of sulphate than sulphide. Siever suggested
that C1 is removed from the atmosphere by freezingout with CO2 and H 2 0 ice, and doubted theexistence of evaporite sediments because of theephemeral nature of liquid water on Mars.
Because of the absence of subduction on Mars,
Siever suggested that extrusive rocks would become drier with time, beginning with early wetbasalts of low viscosity and ending with dry basaltsof higher viscosity.
Perhaps partial melting on Mars might becomplex. For a hydrated silicate composition withhigh Mg/Fe, partial melting (fig. 2) might begin
below IOOOK at a pressure of lO-2O kb. Uprise of
the magma might remove the radioactive elements
making it difficult for the remainder to become
heated high enough for further partial melting. The
key to the outcome is how small a degree of partial
melting is required before (diapiric?) separation
begins. Presence of FeS and Fe metal would cause
further complications, as would a high Fe/Mg ratio
in the silicate fraction. There is a rich field here forstudy. Whatever the outcome Fe-rich lavas are to
be expected on Mars, as is consistent with the high
Fe/A1 ratio in the Viking XRF analyses.
Water- and C02-rich basaltic liquids might retain
most of their volatiles in crystallized minerals (e.g.
amphibole and calcite) rather than releasing themto the atmosphere. Any ash deposits with substantial glass might alter to clay minerals or
zeolites. There is no morphological evidence on
Mars for the extrusion or intrusion of viscous

rhyolitic lavas like those on Earth, but ferrorhyolites might have low enough viscosity to match
high-level flows on Martian volcanoes.
With this geochemical and geophysical background it is not surprising that the inorganic
analyses of fines at the Viking sites appear to be
dominated by weathered Fe-rich silicate and sulphate. Plausible mechanisms can be invented fortermination of volcanism on Mars some lO8 lO9 yrago, especially if Mars melted almost completely
before - 4 x lO9 yr, and it may be possible for Mars
to have incompletely differentiated if H20- andCO2-rich partial melts removed the radioactive
elements from the interior. Finally, it is quitepossible to envisage mineralogical traps for volatiles, thereby allowing a thin atmosphere even forbulk Martian compositions rich in K, H20, CO2,
Nz, and inert gases (Rasool and Le Sergeant, I977;
Fanale and Cannon, 1978; Fanale et al., 1978 ).
Unfortunately it will prove very difficult to placemineralogical limits on the storage capacity for thevolatiles, and hence to decide what types ofmaterial contributed to the accretion of Mars.
Currently it is probably easier to model Mars as avolatile-poor planet (Anders and Owen, I977) , andto explain the mineralogy in terms of a much lower
content of H 2 0 than the one H / O for every two Ca
proposed by Lewis (1972).
Future mineralogical explanation of Mars
should bear in mind the complex weathering
processes, recently considered from the thermodynamic viewpoint by Gooding (1978). Determination of the chemical composition of unweathered
volcanic blocks, preferably of a range of ages,
would be more valuable than study of the finegrained material. If possible, the surface layer
should be spalled off to yield access to unalteredinterior. Return of a soil sample to Earth would
result in much valuable information, but every
effort should be made to design a mission to recoverunaltered volcanic rock, perhaps from the rim of ayoung impact crater, or perhaps from a drill hole.
Phobos and Deimos

The mineralogical nature of these two tiny
irregular moons is unknown. The low albedo andheavy cratering may indicate a regolith (Zellner andCapen, 1974; Veverka, I974; Veverka and Duxbury, I977). If spectral data and the density(c.2 g/cm 3) correctly indicate a composition likea carbonaceous chondrite, these moons areunlikely to be the last survivors of a debris cloudformed by fission or disintegrative capture, and it
is perhaps more profitable to speculate that they
are captured C-type asteroids.

MINERALOGY

OF

THE

PLANETS

55

COMETS, ASTEROIDS, .AND METEORITES
There's not the smallest orb which thou behold'st
But in his motion like an angel sings,
Still quiring to the young-eyed cherubins.
William Shakespeare, The Merchant of Venice, V. i. IO.
Planets underwent so much differentiation that
it is difficult to 'see' back to the solar nebula.
Meteorites yield crucial information on the earlyprocesses of differentiation, but uncertainty about
their original locations hinders speculation aboutchemical differentiation of the condensing nebula.
Spectral studies on asteroids (often called minor
planets) are opening up new vistas, but interpretations and speculations are controversial. Little is
known about the mineralogy of comets. Extremely
important for understanding the mineralogy of theplanets will be landing and fly-by missions toasteroids and comets, though technical problems
will be severe. Until then, most speculations about
the inter-related origin of asteroids, comets, andmeteorites will be difficult to evaluate. A book
Comets, Asteroids, Meteorites, Interrelations,
Evolutions and Origins (Delsemme, 1977), was inpress when this section was written.

Comets and interplanetary dust
Differentiation of the solar nebula probably
involves ejection of much of the mass as either gas
and dust propelled by the solar wind, or as
planetesimals perturbed by the giant planets. [Note.
Small particles could spiral towards the Sun by the
Poynting-Robertson momentum process.] Other
stars would be ejecting material, while condensation would be taking place in the inter-stellar
regions. Ejected planetesimals would undergo
several types of orbital perturbations, and it isqualitatively easy to envisage the Sun being surrounded by a huge cloud of heterogeneous bodies
with orbits ranging from elliptical to parabolic,or even hyperbolic. This Oort cloud has been
estimated to consist of IO6 to IO11 bodies, but
only those bodies are visible that approach the
Sun closely, and have enough volatiles to form
a visible tail. Comets with small elliptical orbits,
produced by planetary perturbations, lose
volatiles during each close approach to the Sun,and should ultimately burn out and become
asteroids if they have a coherent residue. Fromthe mineralogic viewpoint the only 'direct' evidence is that comets have several types of tails.
Those with linear streamers give spectra of ionizedmolecules, including CO +, N ] , CO~, OH +, andionized atoms. Diffuse tails are dominated by thereflection of sunlight, and must consist of dust

grains. Whipple's model of a 'dirty snowball' isgenerally accepted. The proto-comets would
accrete dust particles and ices in the outer solar
system, and some might scavenge gas and condensates in inter-stellar space. Upon approaching
the Sun the ices would become unstable undersunlight.
Dispersed particles are the cause of meteor
showers, and interplanetary dust particles may
result dominantly from short-period comets andonly slightly from collisional debris from asteroids.
Lunar surface rocks are pitted by micrometeorite
craters, but the high impact velocity results inmajor loss of material. Readers should consult theProceedings of the Lunar Sciences Conference fornumerous papers, and should note that lunar soil
contains about I ~o o f C I composition attributable
to meteorite accretion.
Perhaps the best indication of cometary material
is in the fine particles settling into the Earth's
stratosphere (Brownlee et at., I976, I977). Theoretical calculations for spherical particles arriving at
15 km/s indicate that ones greater than IO #m areheated to IOOO ~ or higher for a few seconds
whereas ones below 2 /~m are hardly affected.
Particles collected at 2o km height consist of fivebasic types, namely 'chondritic', Ni-bearing Fesulphide, Fe-poor olivine and pyroxene, N i - F emetal, and others; 6o ~o of the 5- I o ~tm particles are'chondritic' and 3o~o are sulphide, whereas particles >3o #m are dominated by olivine andpyroxene. The particles are more porous and finergrained than those in CI and Cz meteorites, andcommonly are aggregates of grains in the IO-5OO
nm range. Electron microprobe analyses of thirteen
micrometeorites (3-3o ttm) have a two- to ten-foldrange of major elements, with a general similarity
to those of C2 meteorites but with slightly lower Fe,
Mg, Ca, A1, Ni, Cr, and Ti and higher S. Somechondritic aggregates have C and S > 4 wt ~o, andsome contain magnetite and Ni-pyrrhotine as well
as Fe-poor olivine and pyroxene, and Ca, A1,Ti-richmaterial. Together with tentatively identified
hydrated silicates in some particles the assemblageis strongly reminiscent of carbonaceous meteorites.
Deep-sea sediments (Blanchard et al., I978 )contain numerous spheres of three types. Iron-rich
spheres, which contain magnetite accompanied byvarying amounts of wfistite, hematite, taenite, andtrevorite, were identified a century ago, and are

J. V. SMITH

56

attributed to ablation of iron bodies because ofidentical mineralogy with fusion crusts of ironmeteorites. Silicate spheres have etched surfaces
revealing euhedral olivine and magnetite crystals,
but the interiors have surviving glass of various
compositions. Pentlandite, Ni-rich troilite, andchromite also occur. Although most of the silicate
spheres melted completely and then recrystallized
into various textures, including the barred-chondrule type, some contain surviving grains of forsterite, enstatite, Fe-spinel, chromite, pentlandite, andNi-Fe. Bulk electron microprobe analyses of theleast-etched spheres showed a range of compositions (SiO 2 3~ 52; A1203 I 3; FeO I8-45; MgO
I8-27; CaO i-3; NiO t-3) generally describable aschondritic except for low Ni. This second group of
spheres is attributed to ablation of a wide range ofstony bodies. The third group of glassy sphereswith skeletal magnetite may derive from ablation of
metal-rich silicate bodies. Rarity of sulphide may
result from preferential loss by vapourization as
meteoroids hit the Earth's atmosphere.
Perhaps sophisticated mineralogical and microscopic studies will reveal details of crystallizationand irradiation histories, and allow definitive conclusions about the ultimate source of the interplanetary grains. There is an enormous literatureon the composition of interstellar grains (to belocated in astrophysics journals: see obsolete bookedited by Field and Cameron, I973). Bearing in
mind the observational problems coupled withtheoretical uncertainties (e.g. on light scatteringfrom silicate grains coated with ices; competition
between spluttering and condensation), it seems
reasonable to interpret the current observations on
interstellar grains in terms of material similar to or
identical with the interplanetary grains, and to
relate both to fine-grained material found inmeteorites.
Asteroids
General description. The asteroids are difficult to
describe succinctly because of the diversity oforbital elements and surface properties. They canbe separated into three main groups: the Trojan
asteroids, which lie within 2o ~ of the Lagrangianstability points in the orbit of Jupiter; the Earthcrossing and Mars-crossing asteroids; and themain-belt asteroids whose orbits lie between Marsand Jupiter.
Trojan asteroids are too remote for easy spectral
study from Earth. Their low albedo would not be
inconsistent with theoretical expectations that they
are primitive planetesimals containing darkminerals such as magnetite, troilite, and carbon.
Published spectra are unlike any from the mainbelt asteroids, and Chapman (I 976) stated that they

do not match with any known meteorites, including
carbonaceous ones.
Some thousands of main-belt asteroids are
known, of which the largest are Ceres (diameter
I ooo km), Pallas (6 ~o), and Vesta (54o). Their orbits
have a mean eccentricity of o.~5 and a mean
inclination of IO~ which are greater than for theplanets and less than for comets. Their mean
distances from the Sun range from 2.o to 3.3 AU(compared with Mars 1.5 and Jupiter 5.2). Asteroids are erratically distributed over the main belt,
but are absent or rare at or near the Kirkwood
gaps, which correspond to orbital resonance with
Jupiter. The present root-mean-square collisionvelocity of 5 km/s in the asteroid belt is some tentimes the escape velocity of several hundred m/s for
Ceres and Vesta, and much greater than the escapevelocity of the smaller asteroids. Many known
asteroids with similar orbits can be grouped into
Hirayama families.
Dynamical interpretation is difficult, but theideas of several workers were brought together anddeveloped by Chapman (1976), Shoemaker (I 977a),and Wetherill (I977). Much more mass was originally present in the main belt (perhaps C. IO 2 0 i o zv g compared to a present estimate ofc.Io 24 g),
but perturbations from Jupiter increased the encounter velocity and orbital eccentricity so much thatplanetesimals smashed each other to pieces instead
ofaccreting into a planet. Most debris was removed
into the inner and outer parts of the solar system byvarious mechanisms. The present population consists of a few fortunate survivors, which are being
degraded by collisions, some of which are sufficiently catastrophic to yield large fragments
belonging to Hirayama families. Material near the
Kirkwood gaps (especially the 2: r gap) is being
perturbed into eccentric orbits, which ultimately
approach Mars. Further perturbations can lead tocapture by Mars, or even deflection into orbits thatapproach the inner planets or the Sun. In early days
the asteroid belt would have been bombarded byplanetesimals and debris deflected from the outer(and inner) solar system, but calculations of thecollisional frequency and consequences are very
difficult.
All Mars- and Earth-crossing asteroids are small,probably with maximum diameter up to some tenso f k m across (e.g. Eros 13 x 15 x 36 kin). The Apollogroup, with orbits intersecting the Earth's orbit, isof particular interest. Shoemaker (~977b) estimatedthat there are IO3 Apollo asteroids with absolute
visual magnitude of at least i8, and that there are
Io 5 of size IOOm or greater. Only two dozen havebeen detected so far. Estimated impact velocities on
Earth are r5 to 4 ~ km/s with root-mean-squarevelocity 25 km/s, and the present production rate ofM I N E R A L O G Y OF THE P L A N E T S
impact craters IO km across should be about one
per million years, in good agreement with observed
values.
Some Earth- and Mars-crossing asteroids haveelongated orbits similar to those of short-period
comets, and might actually be cores of burnt-out
comets. Are Earth-crossing asteroids derived primarily from comets or from perturbed fragments ofmain-belt asteroids, and are they the main sourceso f meteorites ?
Earth-crossing debris less than c.200 m across isextremely difficult to detect, and the numbered
main-belt population is incomplete for diameters
smaller than 50 kin. The recent determination of
the orbit of a tiny asteroid between Saturn and
Neptune suggests that there may be undiscovered
belts between the giant planets.
Remote-sensing studies. Over the past decadereflectance spectrophotometry, optical polarimetry, and thermal radiometry have been developed for remote sensing of planetary surfaces,
and the results have been checked by direct sampling of the lunar surface. Most studies of spectralreflectance of asteroids were made from c.o.3I.I pro, which wavelength range encompasses
the visible region o.4-o.7 #m. Unfortunately, themineralogical interpretations are uncertain, but
recently the available range was extended deeper
into the infra-red by use of bigger telescopes fittedwith new detectors. The narrow-band study (o.92.7 pm; 28 c m - 1 resolution) of Larson et al. (I976)
on the Apollo asteroid 433 Eros points the way toutilization of the detailed spectral differences ofrock-forming minerals in the infra-red. The present
account is based mainly on the reviews by Chapman, Morrison, and Zellner (I975) , Chapman
U976), Morrison (i977) , and Gaffey and McCord
0977, I978), but is too brief to do justice to thecontroversial details.
Asteroids rotate rapidly with a period of several
hours, and the constancy of relative spectral reflectance and polarization except for Vesta (Degewij
and Zellner, 1978) indicates that there is no major
chemical variation from one side to the other, while
the large variations of total flux for some asteroids
indicate a highly irregular shape. Because of their
low escape velocity, asteroids are being 'sandblasted' by collisions, but some debris must remainon the surface (at least several mm) to explain
optical and radar reflectances. Optical reflectances
are interpreted in terms of light penetrating deeply
into an unconsolidated coating of dust particles
whose chemical composition is representative of
the outer part, but not necessarily of the deep
interior of the asteroid. The particles are 'fresh' andhave not undergone the age-darkening that
occurred for a mature lunar regolith.

57

The simplest classification of asteroids uses thegeometric albedo. Variation of the polarization ofreflected light with phase angle demonstrates that
asteroidal surfaces are dusty, and yields an estimateof the albedo. The absolute brightness then yields
an estimate of the scattering cross section, andhence the mean diameter, which can be checkedfrom thermal infra-red measurements. Zellner et al.
0977a) obtained calibrations by reflection polarimetry of lunar soils and meteorites, and were able
to match asteroidal polarization data only withsamples in which large particles were completely
covered by fine dust. Perhaps electrostatic forces
are important, as for lunar fines. Asteroidal albedos
are strongly bimodal, and were classified into C andS types with albedos near 0.o35 and o.i5, respectively. The albedos of the very dark C type werelower than those of lunar fines and meteorites,
though carbonaceous meteorites were only slightly
brighter (albedo c.o.o5-o.o7). Artificial mixtures of
Mg-silicates with IO% carbon were similar in
darkness to C types, and Zellner et al. (I977b)
suggested that the interplanetary material studied
by Brownlee et al. (I977) might be related to C
asteroids, thereby implying a 'supercarbonaceouschondrite' composition.
Spectrophotometry in the visible and near infrared regions offers the best clues to the mineralogy.
Laboratory experiments with mineral mixtures
showed that the spectral signature is relatively
unaffected by grain size so long as the grains have
similar reflectances, and the individual grains have
low absorbance. Gross dissimilarity of reflectance,
as between silicate and metal, causes serious problems for interpretation of spectra from mixtures, as
does large grain size (c. I cm) of silicates. Telescopic
studies of reflectauce of the Moon have shown
good correlation with mineralogical identificationof returned fines, and for the time being it is easiest
to assume that asteroidal surfaces are covered bywell-mixed fine material.
Two approaches have been used for interpretation of spectral signatures from asteroids: directcomparison with laboratory spectra for powdered
meteorites measured by Gaffey (1976), andattempted 'reconstruction' of the observed spectrum from those for individual minerals that areabundant in meteorites (Adams, 1975).
The first approach (numerous studies reviewed
by Gaffey and McCord, I977) was particularly
successful for the asteroid Vesta, whose spectrum
matches those for basaltic achondrites, especiallywith respect to absorption bands at o.92 #m andI. 9 #m indicative of pyroxene with Io-12 mole%
Ca. Quite good matches were obtained between
other meteorites and asteroids (e.g. Chapman,i976 , Fig. I), but many asteroidal spectra could

58

J. V. SMITH

not be matched. Indeed Gaffey and McCord(I978) questioned the validity of some of theearlier apparent matches, and emphasized thedesirability of the reconstruction procedure.
Briefly, the matching approach led Chapman(1976) to conclude that the darker asteroids (C type)may be subdivided into ones like C2 meteorites andones (called C*) perhaps similar to the metamorphosed C4 Karoonda meteorite; that theredder and brighter asteroids (S type) mostly
cannot be matched with meteorites, but may consist of a mixture of silicates and iron in the ratio z: 1
to 1:4, probably closer to stony-iron meteorites
than ordinary chondrites; that Vesta matches withbasaltic achondrites, perhaps most closely witheucrites; that a few asteroids (M type) match
enstatite chondrites, or just possibly iron meteorites or enstatite-iron mixtures; that approximately
forty spectral sub-types can be tentatively recognized (e.g. the fourteen listed in his Fig. 4); and thatmost meteorite types are not matched by thepresently studied asteroids.
Gaffey and McCord (1978) were unhappy with
Chapman's use of the words Carbonaceous andSilicaceous or Stony-iron as the sources of thedesignations C and S, because the important spectral features of a C type might match with a blackchondrite or carbon-poor C4 meteorite and notnecessarily with a carbon-rich meteorite. Nevertheless they agreed that the C-S classification was
useful for reconnaissance, in which a simple
measurement of colour or polarization is made.
The important spectral features for common
minerals in the solar system are: Ices have highalbedos like the Martian ice caps, and are notpossible for asteroids. Feldspar has no absorption
feature in the o.3-I.1 ~tm range, but has a weakabsorption near 1.25 /tm if Fe z+ is a minor
substituent, as in lunar volcanic rocks. Iron-richmetal has no absorption feature, but the generaltrend of reflectance with frequency is more linear
than for silicate. The spectral signature may varywith Ni content. All silicates containing abundanttransition-metals have a strong increase of absorption in the blue and ultra-violet range caused bycharge-transfer. Pyroxenes give a range of spectra:
Fe-bearing varieties have a symmetrical absorption
near o.95 pro, whose position depends on the
Ca/Mg/Fe ratios; Fe-free varieties have a flat andfeatureless spectrum; and Ti-rich specimens have astrong absorption near o.6/~m. Olivines also give a
range: Fe-rich varieties have a broad, non-symmetrical absorption near I #m, whereas Fe-free
forsterite has a flat and featureless spectrum.
Hydroxylated and hydrated silicates (clay minerals)
have a weak absorption near o.6/~m and a strong
absorption near 3 /~m. Hydroxylated iron oxides

have strong absorption near 0.7/tm. Graphite andcarbon-rich compounds have a flat, featureless.
spectrum with very low albedo. Spectra frommixtures are determined principally by the darker
components, e.g. relatively transparent olivine isdifficult to detect in the presence of relatively
opaque pyroxene, while graphite and other very
dark materials such as magnetite strongly reduce
spectral features from intermixed silicates. The
finer is the grain size of a darker phase, the greater is
its dominance in spectra from mixtures with brightphases. So far, the important phase troilite has notbeen considered in the interpretation of asteroid
spectra, but it may be a significant opaque phase on
asteroids, especially if it is fine-grained, or if it iscoating other minerals.
Originally McCord and Chapman (I975)
emphasized the following parameters for asteroid
spectra: R/B, reflectance ratio at red (0.7/~m) andblue (o.4 #m), which tends to separate C from S
types; BEND, curvature, which tends to measuresilicates/(opaques and metal); DEPTH, depth ofpyroxene band at o.95 #m; and IR, reflectance at
1.o5/~m minus that at o.73 #m, negative values ofwhich indicate olivine. The difficulty of detailed
interpretation is pointed up by the controversy
over the different types of data for Eros (Larson etal., 1976; Pieters et al., 1976).
Gaffey and McCord (1978) classified asteroid
spectra as follows: The R A type has near-linear
decrease o f reflectance vs. frequency for visible light(hence Fe, Ni-metal component) and a silicate
absorption feature near I #m (hence olivine or
pyroxene or both). Sub-groups R A - I and RA-2
were chosen depending whether olivine or
pyroxene is indicated as the dominant silicate.
Many subtle differences were discussed including
the Ca content of the pyroxene, the Ni content of
the iron, the presence of plagioclase, and possible
presence of a wfistite coating. The R R type has
near-linear decrease of reflectance vs. frequency,
but no detectable silicate absorption (hence Fe,Nimetal component and possible Fe-free silicate).
Low albedo of some R R asteroids might result
from a fine-grained opaque phase. There is apossible weak absorption near o.6-o.7/~m. The R F
type has decreasing reflectance with frequency inthe visible range and near-constant reflectance in
the infra-red (hence dominated by metal plus Fefree silicate?, but Ni-content of metal and possible
presence of opaque phases such as carbon ormagnetite is uncertain). Perhaps troilite will prove
important in R R and R F spectra. The A type for
Vesta is indicative of basaltic achondrite (eucrite?),and for Dembowska is indicative of olivine-richachondrite with possible minor pyroxene andmetal. The F type has a flat spectrum from o.4-1 #m

M I N E R A L O G Y OF THE PLANETS(and low albedo 0.02-0.05); interpretation is very
difficult, but Gaffey and McCord preferred a
mixture of Fe-bearing silicate (olivine?) andmagnetite, while not ruling out a mixture of opaquephases and Fe-poor clay minerals; a definitive
interpretation is needed since the large asteroids
Ceres and Pallas belong to this group. The T type(for transitional between F and R) has a fairly fiat
spectrum with weak absorption at o.6-o.7 #m and a
UV absorption feature; most C-type asteroids
belong to this group, and some undetermined
mixture of opaque phase(s) and Fe 2 +-Fe 3 +-silicate(s) is indicated. Sub-types TA, TB, and TC were
proposed with approximate correspondence with
C1 and C2 meteorites. Sub-types TD and TE withintermediate albedo (9-15 ~o) and a significant UVfeature indicative ofmafic silicate (but no infra-red
features) pose serious identification problems,though magnetite and/or carbon-rich material (ortroilite?) are indicated.
The Gaffey-McCord approach provides the following possible matches between main-belt
asteroids and meteorites (1977, Table 1): A, eucriteand olivine achondrite; F, C4 (Karoonda); RA,various stony-irons, including mesosiderite, pallasite, and olivine- or pyroxene-rich stony-irons; RF,enstatite chondrite or iron; and T, C1, C2, and C3.
ZeUner et al. (I977) classified 44 Nysa, 64 Angelina, and 434 Hungaria as E-type asteroids, whose
spectra indicate Fe-free silicate and a plausible
match with enstatite achondrite. Nysa is highly
aspherical, and its dynamic association with 135
Hertha and several small objects in a Hirayama
family suggested fragmentation ofa percursor with
Hertha as core and Nysa as a mantle fragment.
Data for Earth- and Mars-crossing asteroids aresparse. Zellner and Bowell (I 977) found only one Ctype object)n a dozen members of the Apollo andAmor groups. Particularly important is the possible matching of Eros (Zellner, 1976) and Torowith high-grade chondrites (H or L; grade 5-6?)
and Alinda with C3.
Spatial distributions and implications. For thesimple S - C - M - E - U classification [U--unclassifiable], there is a strong tendency for increasing
abundance ratio (S/C) with decreasing distancefrom the Sun of main-belt asteroids (Chapman,
I976; Zellner and Bowell, 1977), and this trend
apparently carries inwards to the Apollo and Arnor
asteroids (though the statistics are poor). An earlysuggestion that S objects predominated over C
objects for asteroids with c.15o km diameter wasapparently the result merely of statistical bias.
Gaffey and McCord (I977, Fig. IO) concluded that
the C2 and C3 types (assumed to be primitive
condensates, not heated significantly) occur
throughout the main belt, whereas C4 and differ-59

entiated metal silicate types tend to occur mainlyin the inner half of the main belt. Details are
disputable (especially with regard to the C 4 type),
but the over-all spatial trend of C/S or primitive/differentiated appears plausible.
The implications of the observations are unclear
in view of uncertainty in the evolution of asteroidal
orbits and the extent of catastrophic fragmentation(refs. in Wetherill, I977; Shoemaker, 1977a, b). The
main-belt asteroids could have formed in place or
have been perturbed into the main-belt either fromthe inner or outer solar system. Of course, thesimplest idea is that C- or primitive asteroids werethe 'primordial inhabitants' produced by primarycondensation , whereas S- or differentiated asteroids were perturbed into the main belt from theinner solar system. Attempts to reconstruct theprecursor asteroids from presumed collisional
fragments in Hirayama families (e.g. Gradie andZellner, 1977; Zellner et al., I977) must be regarded as tentative, though fascinating. The
sources of heat for differentiated asteroids are unknown, though plausible possibilities exist (Sonettet al., 197o), and it is not known when
the heating occurred, except by implication frommeteorites. Gaffey and McCord (I977) proposed that all the larger asteroids apparently
have surfaces of differentiated material (either silicate, metal, or C4 type), but whether this results
merely from greater mechanical strength andresistance to disintegration of heated asteroids orsome correlation between heat sources and size is
not clear. The recent discovery of water ofhydration on Ceres, the largest asteroid, by Lebofsky (I978) is inconsistent with this general proposition, and may require some revision of the spectralinterpretation of C-type asteroids.
Although there are possible spectral matches
between asteroidal surfaces and most types of
meteorites, it is far from clear how meteorites
derive from asteroids. Current dynamical theories(refs. in Wetherill, 1977) utilize the concept of
perturbation of collision fragments from asteroids
near the Kirkwood gaps thereby ultimately providing Earth-crossing material. Actually the spectral
characteristics of Apollo asteroids appear to be
dominated by S-type material in contrast to C-type
dominance in the main belt. Perhaps the primitive
main-belt aSteroids yield finer debris than differentiated asteroids, and perhaps further study ofApollo asteroids will yield a correlation between
size and spectral type. Particularly problematicalfrom the dynamic viewpoint is attempted assignation of the supposed eucrite parent planet to theasteroid Vesta, which is far from an orbital resonance with Jupiter, but see Consolmagno and Drake(1977) and Hostetler and Drake (1978). Readers are6o

J. V. S M I T H

referred to Chapman (1976) and Gaffey and
McCord (1977, 1978) for provocative speculations
about the origin and development of asteroids, andto O'Leary (I977) for a proposal to mine Apollo
and Amor asteroids, which just might enhancefunding for scientific exploration.
Meteorites

All planetary models are explicitly or implicitly
based on analogies with meteorites. The 'Heroic
Age' of meteoritics (i.e. before the Apollo I I
mission) produced valuable but over-simplified
ideas, and recent petrographic data interpreted in
the context of solid-state theories are leading tocomplex interpretations involving condensation,accretion, crystal liquid differentiation, collisiona1
destruction, and metamorphism. For brevity, thefollowing sections are rather selective and didactic.
Recent general reviews are: Buchwald (I 975), Nagy
(1975) , Wasson (I974) , and Mason (1971b). Older
but still valuable are: Keil (I969a), Anders (I964),
and Mason (1962) .
Carbonaceous meteorites. Three dozen silicaterich meteorites carry substantial C and werecharacterized from bulk chemistry into Wiik types
I, II, and III (data from Mason, I971b):
Wt~o
CI
CII
CIII

Si

A1

Fe

Mg

Ca

Na

C

S

H20

IO
I3
I6

0.8
I.I
1.4

I8
22
25

IO
I2
15

i.I
1.3
1.7

0. 5
0.4
0.4

3
2
0.5

6
3
2

20
13
I

All carbonaceous meteorites are mechanical mixtures of materials that have reacted with gases or
liquids at diverse temperatures. They coalesced at
low temperature, and are grossly out of mineralogical equilibrium. A petrologic classification into
CI, C2, C3, and C4 types depends on ill-definedfeatures indicative of increasing consolidation, anddetailed study is leading to sub-division of thecrude chemical classification.
Five CI meteorites (typified by Orgueil) havebulk compositions whose non-volatile elements fit
closely with the composition of the solar photosphere (Table II). They are regarded as the bestindicators of the chemical composition of the solar
nebula when appropriate amounts of H, rare gases,
N, and C are added. However, they could not haveformed as direct equilibrium condensates from thesolar nebula, and the bulk chemical composition
results from addition of many components. The CI
meteorites are friable, and hydration of sulphates
may have occurred since landing on Earth. Nagy
(I975) describes the many mineralogical studies
and lists twenty-five minerals. Phyllosilicates, too
fine-grained for easy study, dominate the matrix.
Bass (197 I) identified montmorillonite and serpen-tine, and Kerridge (I976, 1977b) and McSween andRichardson (1977) interpreted bulk electron-microprobe analyses of fine-grained matrix in terms of
Mg-rich phyllosilicate(s) and a NiS-bearing component, perhaps the complex product of aqueousalteration. The carbon occurs mainly as a wide
variety of organic compounds (Nagy, I975), whichhave been interpreted by Anders and co-workers as
the products of abiogenic synthesis. Carbon also
occurs as several carbonates (Fe-bearing magnesite, calcite, dolomite). Magnetite occurs in several
morphologies. Rare olivine and pyroxene grains
have wide composition ranges (Reid et al., 197o),
and Kerridge and McDougall (1977) discovered
particle tracks in only 15 ~ of the olivine grains.
Sulphur occurs as very rare troilite, pyrrhotine, andpentlandite, as native sulphur, and mainly assulphates in complex veins: hydrated Mg- and Casulphates occur, but the extent of terrestrial hydration is uncertain. There are no chondrules in CI
meteorites. Mechanical aggregation of diverse
material on a disturbed planetesimal regolith,
together with aqueous alteration (Du Fresne andAnders, I962), provides a plausible origin.
The remaining members of the carbonaceouschondrites actually do contain chondrules, thoughfine-grained matrix is dominant. Nomenclaturedepends on features selected by the classifier(Mason, 197Ia; Van Schmus and Hayes, I974;
Wasson, I974, 1977; Fitzgerald and Jones, 1977).
Type CIII (c.C3) was subdivided into Ornans (O)and Vigarano (V) varieties (Van Schmus, 1969), thelatter including the large Allende shower. Type CII(c.C2) includes the Murchison, Murray, and Mighei(M) meteorites. All these meteorites are mechanicalmixtures of several components, though the bulk
chemical compositions (Table II) can be approximated by simple mixing models related to equilibrium condensation theory. The Karoonda (C4)meteorite is interpreted as the product of opensystem metamorphism of a C3 meteorite on thebasis of volatile-element depletion (Matza andLipschutz, 1977). Chemical and mineralogical properties are too complex for detailed description
here. When normalized to Si, the abundances of Ca,
A1, and Ti are lower and of Fe and Ca are higher in
CI and most C2 and C3(O) meteorites than in
C3(V) and a few C2, C3, and C4 ones. Volatileelements tend to correlate inversely with the abundance of high-temperature minerals, and the bulk
composition can be matched with equilibriumcondensation theory (e.g. the'step' pattern in C3(O)and C3(V) meteorites), but with the possible complication of dust-gas fractionation in the Ornans
sub-type (Anders et al., 1976). But in detail thehistory of these meteorites cannot be interpreted
in terms of equilibrium condensation, and inM I N E R A L O G Y OF THE PLANETS
particular the oxygen isotopes require mixingof unrelated components.
The complex features of CM meteorites areyielding to modern techniques, but the observations and interpretations are incomplete and somewhat controversial. Detailed studies of Murchison(Fuchs et al., I973) and of Murchison, Murray, andclasts in the Jodzie howardite (Bunch and Chang,
I978a, b; Bunch et al., I978) provided most of thefollowing description. Textural relations, which
include schlieren, foliation, and boudinage structures, indicate that the CM meteorites once resided
in a regolith breccia. Pseudomorphing of clasts byphyllosilicates and poorly characterized phases
indicates subsequent alteration. Murray contains
xenoliths mainly of CM type, but Io ~ of C3 type inwhich phyllosilicates replaced olivine andpyroxene. Chemical and structural identification of
the phyllositicates in the CM meteorites is difficult
and Bunch and Chang (I978b) found at least fourtypes in Murray, probably corresponding to either
septechlorites or serpentines. Electron microprobe
analyses showed considerable C, S, and Ni, which
probably result from intergrown submicroscopic
phases. Opaque materials for several textural varieties, collectively described as poorly characterizedphases, consist of a mixture of an Fe-rich phyllosilicate and a phase enriched in Fe, Ni, C, and S (Bunch
and Chang, 1978b ). These authors attributed these
opaque materials to alteration of breccia matrix,
rather than to direct nebular condensation(McSween and Richardson, i977). Of the sulphides,
pyrrhotine and penflandite are texturally early, as
is a mysterious phase that probably resembles thecarrier of a strange xenon component in Allende(Lewis et al., 1975). Olivine grains comprise IO3O~o of CM meteorites (Mason, i963a; Wood,
I967), mostly with o - 5 ~ Fa, but up to 8 o ~ Fain Murchison with Fe and Ni positively correlated(Fuchs et al., 1973). Direct condensation of theolivines from the solar nebula was favoured by
Grossman and Olsen (1974) and Olsen et al. (~977)
because of presence of Ni,Cr,P-rich metal inclusions. Attribution of the olivine to release fromchondrules upon conversion of glassy mesostasis tofriable phyllosilicate (Richardson and McSween,
I978 ) is apparently negated by the observation of
C-rich mounds on the surface of the olivine grains(Bunch and Chang, 1978b). The occurrence of rarehibonite and perovskite (high-temperatureminerals) with gypsum and whewellite (lowtemperature minerals) reinforces the other evidence for mineralogical disequilibrium. Compaction of olivine and phyllosilicate at -4.42 x l O 9 yr
for Murray and - 4.22 X IO 9 yr for Nogoya is indicated by fission-track densities (MacDougall
and Kothari, I976).

6i

For the C3 meteorites, McSween (I977b, c)emphasized the petrographic diversity. The
Vigarano subgroup has larger chondrules andgreater abundance of dark matrix than the Ornans
subgroup, as well as higher Ca, A1, and Ti and lower
Fe. Of the fourteen Vigarano meteorites, two(Renazzo and A1 Rais) have phyllosilicate matrices
and two (Coolidge, Mulga West) have been metamorphosed to Van-Schmus-Wood grades 4, 5, or 6.
The main group of ten meteorites (including
Allende) has 32-51 ~ by volume of Type I chondrules (granular forsterite, Fao_lo, clinoenstatite,
Fso_7, Ca-Al-rich glass; some with no opaqueminerals, but most riddled with metal, magnetite,
and/or sulphides). There are traces of Type II
chondrules (olivine, Fa2o-5o, and Ca,Al-rich glass)and Type IV chondrules (melilite, anorthite, spinel,
diopside, glass), and four types of melilite-richchondrules and aggregates were distinguished byMartin and Mason (I974). Amoeboid olivine(Fa1_36) aggregates with minor pyroxene (fourtypes), spinel, anorthite, perovskite, nepheline, andsodalite (Grossman and Steele, I976 ) amount to I9 ~o. Two types of Ca,Al-rich inclusions [Type A
8o-85 ~ melilite, 15-20 ~oo spinel, I-2 ~o perovskite,rare anorthite, hibonite, wollastonite, grossular,
pyroxene (< 9~o A1203, < o.7~ Ti); Type B 356 o ~ pyroxene (> i 5 ~ A1203, > 1.8~o Ti), 153O~o spinel, 5-25 ~ anorthite, 5-2O~o melilite]amount to 3-9 ~, but see Blander and Fuchs (I 975)
for a four-fold classification and arguments forcrystallization from a liquid rather than directcondensation at high temperature from a gas; ofcourse, melting could have occurred after condensation, as also suggested by Grossman (1975).
Rock and mineral fragments (1-3 ~) include isolated olivine grains, which were attributed torelease from chondrules (McSween, I977a), but see
preceding paragraph. The matrix (35-51 ~) consists mainly of olivine grains, free from radiation
damage (Green et al., x970 and coated and interspersed with carbonaceous material (graphitic?) of
uncertain nature (Bauman et al., I973), and a few
per cent of opaque minerals including metal,
sulphide, and magnetite.
McSween subdivided the Vigarano group into
three groups: five reduced meteorites with 0.5- 5
kamacite plus taenite, 1.5-2.6 ~o troilite, and o.I0.6 ~o magnetite; five oxidized meteorites, including
Allende, with 0-0.2 ~ metal, 1.6-4 ~o Ni-rich pentlandite plus minor troilite, and 1.2-3.2 ~o magnetite; and a reduced pair (Renazzo, A1 Rais) classifiable as petrographic type C2. Perhaps magnetite
resulted from oxidation of sulphide (Herndon etal., 1976). High Co contents ofkamacite fall outside
the main range of meteoritic metal (Fuchs andOlsen, I973) and perhaps require crystallization in62

J. V. SMITH

an olivine-differentiated environment (Smith andSteele, i975).
Several Vigarano-type meteorites, including
Allende, are agglomerates, and many types ofclasts
await detailed study. Fe-rich rims on fine-grained
inclusions in Allende indicate metamorphic transfer of Fe from the matrix, and coarse-grained
inclusions have Na-rich rims and complex texturalfeatures (Hutcheon et al., I978). It is not clear
whether chemical changes occurred prior to or
after final consolidation, or both. Certainly the
Vigarano-type meteorites have some chemical andmineralogical features attributable to high-temperature condensates (e.g. Grossman and Clark,
1973; Grossman and Ganapathy, 1977; Grossman
et al., 1977) but petrographic data require complex
sequential processes, as advocated by Fuchs,
Grossman, and others (e.g. Kurat et al., ~975;
McSween, I977d ). Oxygen isotope data (see later)
require survival of pre-solar grains enriched in 160(Clayton et al., 1977), and a mass of data on isotope
anomalies in Allende inclusions (Clayton, I978) is
pointing to a pre-solar history probably associatedwith a supernova. Trace-element data require complex interpretations (Chou et al., 1976; Takahashi
et al., I978), whose correctness will require testingby sophisticated chemical studies of individual
fractions selected with petrographic control.
The Ornans-type meteorites contain olivine-richchondrules, inclusions (some Ca, Al-rich), rock fragments, matrix, metal, and sulphide. McSween(I977b) placed them in a metamorphic sequence
from Kainsaz to Warrenton and perhaps
Karoonda, in which the olivine and pyroxene
become richer in Fe while the metal loses Cr andgains Co. The ratio of matrix to chondruleschanges.
A tentative new group of carbonaceous chondrites (denoted CK) consists of the Kakangari andAdelaide chondrites, together with the lunar fragment Bench Crater (references in Fitzgerald andJones, I977). All three have chondrules in a magnetite-bearing unequilibrated matrix, and the principal bulk chemical feature is low Ca/A1 (c.o.5
atomic) perhaps resulting from condensation of agas that had lost Ca-rich materials. Some xenoliths
in meteorites may belong to this group. Data for
Kakangari are given later.
Many features of the carbonaceous meteorites
would have been obliterated by metamorphism attemperatures exceeding 5oo-Iooo K. There isabundant evidence of metasomatism in CI and CII
meteorites probably associated with water or
steam. Coexistence of isotopically anomalous
Ca,Al-rich inclusions with olivine-rich matrix
in CIII meteorites requires survival of presolar
grains.

Ordinary chondrites. The ordinary chondrites
have similar chemical compositions (Table II), but
can be separated into three groups on the basis of
total Fe: high (H) c.28 wt ~o, low (L) c.22 and verylow (LL) c.2o. They contain rounded bodies(chondrules), and display petrographic properties
ranging from the Van-Schmus-Wood type 3(sharply defined chondrules; much opaque matrix
and glass; compositionally diverse olivines andpyroxenes; no feldspar and taenite) to type 6(indistinct chondrules; recrystallized matrix and no
glass; uniform olivines and pyroxenes; distinct
feldspar and taenite). In the Heroic Age it waspossible to envisage accretion of chondrules andfine debris from the solar nebula into threeplanetesimals each of near-uniform composition,
followed by metamorphism varying in intensitywith depth. Unfortunately this attractive story (e.g.
Wood, 1962) is insufficient, and it is necessary toconsider six processes for each chemical group:
formation of a debris cloud of diverse components
during a major collision of planetesimals; mechanical sorting of the cloud under the influence of acentral gravitational field; sequential accretion
with change of composition, especially of volatileelements; thermal metamorphism varying withdepth; formation of a regolith, yielding mainly
monomict breccias with varying content ofadsorbed solar gases, but producing some polymict
breccias incorporating late-accreting material; anddisruption of the regolith and underlying metamorphosed material during the collision thatstarted fragments on their paths towards the Earth.
Many petrographic features can be interpretedby analogy with lunar breccias (Prinz et al., I977a).
Chondrites composed of light- and dark-textured
components with the same bulk composition (e.g.
Frederiksson and Keil, t963; Bunch and St6ffler,1974) but different contents of solar-implanted
gases can be explained by impact mixing of aregolith. Brecciated chondrites are mostly monomict, indicating a uniform regolith on the parent
planetesimal, but rare specimens contain inclusions of a different type. Wilkening and Clayton(1974) summarized existing data on inclusions ofcarbonaceous type in brecciated ordinary chondrites, and further examples were given by Fodor
and Keil (I976), Fodor et al. (I976), and Wilkening(I976). Complex regolith annealing and disruption
are implied by occurrence of unequilibrated inclusions in a more equilibrated host in the Weston
brecciated chondrite (Noonan and Nelen, I976).
Strong shock caused melting in irregular veins ofsome chondrites, with formation of disseminatedsulphide droplets and veinlets (e.g. Wahl, I963;
Bergemann and Wlotzka, I969), and severe shockcaused conversion of plagioclase, olivine, andM I N E R A L O G Y OF THE PLANETS
pyroxene to maskelynite, ringwoodite, and majorite in some chondrites (Tenham, Coorara, etc.), andemplacement of melt in the St. Mesmin chondrite(Dodd and Jarosewich, i976 ). Weaker shock
resulted in lithification of some gas-rich chondrites(Ashworth and Barber, i976 ).
Most ordinary chondrites have textures essentially unaffected by impact, and they can be classified into the Van-Schmus-Wood types 3 to 6. Therecan be no doubt that types 4 to 6 have undergone
prolonged thermal metamorphism, and attention is
directed first to the unequilibrated ordinary chondrites of type 3, which retain textures from theoriginal aggregation. Such type 3 chondrites consist of two types of chondrules set in a fine-grainedcomplex matrix. 'Droplet' chondrules have textures(excentroradial; barred) definitely resulting fromsupercooling of spheres of silicate-rich liquid (e.g.
Kurat, I967a; Blander et al., I976), while 'lithic'chondrules have textures matching those of
igneous rocks (e.g. Kurat, i967a; Dodd, 1978). The
early literature did not always distinguish between
the two types, but the later literature recordedfeatures indicative of their origin: e.g. Dodd (i97 I,r974) noted that droplet chondrules were unshocked whereas lithic ones were shocked in the
Sharps and Hallingeberg chondrites. Olivines andpyroxenes have a range of Mg/Fe ratios and diverse
CaO contents indicative of low to high crystallization temperatures (e.g. Dodd, i973). These features
suggest derivation of lithic chondrules by mechanical abrasion of a variety of igneous and metamorphic rocks. The fine-grained matrix contains
volatile-rich material not in chemical equilibrium,
such as the nepheline-rich and silica-oversaturated
components of the Mezr-Madaras brecciatedchondrite (Kurat, 1967b; see also Binns, I968; Van
Schmus, I967). The Bovedy L3 chondrite contains
Anss glass (Graham et al., I976) whereas mostchondritic glasses have high Na/Ca. Detailed
microscopy of the Tieschitz (H3) chondrite (Christophe Michel-Levy, I976) revealed a fine coatingaround the chondrules, an interstitial deposit of
Si,Al-rich material, primary cavities in chondrules
and secondary ones produced by leaching of glass,and, finally, debris of broken chondrules. Clastic
matrices of high porosity were observed in L3 andLL 3 chondrites (Ashworth, I977) as well as nonclastic matrices attributed to solid-state recrystallization.
The chondrules have roughly the same size
distribution in all ordinary chondrites (Dodd,I969). Controversy over the details probably
results from selective disintegration of some chondrules, and the data of Dodd (I976) and Hughes(I977, I978 ) obey a log-normal distribution.
Hughes (I977) noted that the mass distribution

63

matches that in active meteor streams, and Dodd(i976) found that metal-rich particles were smaller
than silica-rich ones. Dodd (x969) described evidence for mechanical sorting during accretion, andin I976 proposed that this resulted during laminarflow of a particle cloud around a growing planetesimal, using aerodynamic ideas of F. L. Whipple.
Such a cloud might result from disintegrative
collision of two planetesimals, or disintegrative
capture of one planetesimal by another.
Evidence from lunar rocks does not favour
generation of ordinary chondrites by minor
impacts into a regolith, because of the preponderance of agglutinates over chondrules in such a
process (Kerridge and Kieffer, I977). Volcanic fire
fountains (Ringwood, 196Ia, I966a) could n o t
produce bodies large enough to give chondrites
with a large variation of metamorphic grade.
Impact of dust grains (Lange and Larimer, I973)
and direct condensation of liquid droplets from thesolar nebula (Wood, i963, McSween, i977d ) do
not provide a plausible explanation of the association of droplet and lithic chondrules in ordinarychondrites if the lithic chondrules form bymechanical abrasion of rocks. However, theseother processes may well have occurred in the preplanetesimal stage, and direct condensation
followed by aggregation may be responsible forporous aggregates in carbonaceous meteorites (e.g.
Grossman and Steele, i976; Wood and McSween,I978).
Another process for formation of chondrules ismelting upon high-speed impact into a planetaryatmosphere, but this also fails to give the association of droplet chondrules with a range of lithic
chondrules.
Early mineralogic studies of metamorphism (reviewed in Dodd, I969) have been extended by many
detailed studies with the electron microprobe.
Temperature estimates based on Mg/Fe/Ca contents of coexisting pyroxenes and olivines are
somewhat uncertain but suggest c.9oo K for type 4
to c.~3oo K for type 6. Oligoclase develops as
distinct grains in type 6, and the olivines andpyroxenes have narrow composition ranges intypes 5 and 6. The Shaw chondrite was placed in
type 7 because of evidence for anatexis and highmetamorphic temperatures near 155o K (Doddet
al., I975). Partial melting of metal and sulphidefrom the Shaw chondrite should have yielded acomposition similar to but not the same as the San
Cristobal IAB iron (Rambaldi and Larimer, I976 ).
Chromite increases five-fold from type 3 to type 6,
and Cr decreases while A1 and Ti increase (Bunch etal., I967). Urey and Mayeda (I959) interpreted themetal particles of chondrites as fragments of a
disrupted planetesimal that had cooled slowlyJ. v. SMITH

64

before accumulation into a secondary object andreheating to 770 K. From zoning profiles of clear
taenite in twelve chondrites, Taylor and Heymann(I 97 I) deduced cooling rates of I to IO K / I o 6 yr andcrystallization temperatures of 61o-66o K while
Smith and Goldstein (1977) deduced very rapid
cooling rates of I and IOO K/day for two severely
reheated chondrites. An unresolved problem is the
deduction by Pellas et al. (1978) that the Shawchondrite must have cooled very slowly to retain
fission tracks whereas metallographic data by
Berkley et al. indicate rapid cooling. Gallium, Ge,
and W increase in metal with petrologic type (Chouand Cohen, I973; Rambaldi, I976), perhaps because of extraction from silicate.
Although the major elements (especially Fe) of
the ordinary chondrites allow classification into the
H, L, and LL groups, irrespective of the details of
brecciation or petrologic type, the minor and traceelements vary with petrologic type: specifically,
Binz et al. (1976) found that Bi, In, and T1 were tenfold lower in equilibrated than in unequilibrated
ordinary chondrites, and attributed this to preaccretion processes in the nebula rather than tothermal metamorphism. Much more study isneeded to sort out the effects of pre-accretionchemical variation and post-accretion metamorphic changes for which the heating studies of
Ikramuddin et al. (1977 a, b) are providing valuable
controls. In the Heroic Age the bulk chemistry of
the chondrites was interpreted as the mixing of ahigh-temperature condensate with a low-temperature condensate, as detailed by Keays et al. (I971).
Actually the success of this model must rely on
statistical averaging of several materials into eachtype of supposed condensate, and planetesimalformation must intervene between nebular condensation and chondrule formation if the Dodd
model is sustained. The chemical analyses of traceelements in individual chondrules by Osborn et al.
(I973), and in separated minerals by Rambaldi,

point the way to further understanding of t h e .
genesis of ordinary chondrites.
It will be assumed now that the ordinary chondrites are the result of accretion of dust clouds
produced by three separate collisions. Perhaps theordinary chondrites represent only the outer partsof accreted bodies and the inner parts containcoarse breccias of blocks too heavy to go into thedust clouds: of course, melting might have caused
major crystal-liquid differentiation of suchmaterial. Such a brecciated interior might be
represented by the Netscha~vo iron (Bild andWasson, 1977), which contains z 5 ~ angular blocks
of silicate with uniformly distributed metal.
Perhaps this silicate-rich brecciated iron meteorite
might be the product of collision of a chondritic
planetesimal with an iron planetesimal. Anyway
the silicate material is less Fe-rich than the Hchondrites, and may point the way to discovery ofother types of ordinary chondrites.
Enstatite chondrites and achondrites: reducedirons. Whereas the CI and CM meteorites haveessentially no Fe in metal and sulphide, and theordinary chondrites have Fe distributed between
metal-sulphide and silicate (H c.3/4 of Fe in metalsulphide; L c. 1/2; LL c. I/3), the enstatite chondrites
and achondrites (Table II) have no or very little Fein the silicates. Indeed the enstatite meteorites are
so reduced that metal contains Si (Ringwood,
I96Ib) and troilite contains Ti (Keil, 1969b), butwith variable partitioning with silicate (Easton andHey, 1968). Also the minerals oldhamite (CaS),
sinoite (Si2N20), osbornite (TIN) occur. The mineral assemblages and compositions match thosepredicted by Larimer (1975) for a nebula of cosmic composition with C/O ~> I rather than the o.6
used to explain the mineralogy of ordinary chondrites and carbonaceous meteorites (Table III).
Most enstatite chondrites (Mason, I966 ) can beroughly classified into two main chemical andmineralogical groups (Keil, I968):

Plagioclase*chondrules
III

volatiles+ to + + solar abundance-depletedFe

Si/Mg S

30-33 1.4
2I 29 1.2

6
3-4

Ab

An

Or

93
2-3
I-4
80 83 13 I7 3 4

kama- troilcite
ite
Si
Ti

oldhamite
Mg

Mn

CaO

MnO

FeO

3-5
t.3

1.4
0.5

0.3
I.I

o.z
0.7

o.~
o

0.6
o.z

0.3-o.6o.6-o.7

enstatite* M o l ~ ; the rest wt%.

Type I belongs to petrologic grade 4, and type II togrades 5 and 6, but they cannot be related byisochemical metamorphism (Keil, I968; Shaw,
1974). Detailed textural studies are needed tocomplement the careful mineralogical data. Type I
chondrites have several types of sharply definedchondrules plus angular fragments (especially of
metal) set in a diverse matrix. Type II chondrites arehighly recrystallized with rarity or absence ofchondrules. Presence of silica polymorphs (mostly
tridymite but some quartz in II) demonstrates ahighly undersaturated composition, while djerfisherite (K-bearing Fe sulphide) shows that K is
chalcophilic under highly reduced conditions. The
sulphides (Ca, Mg, Mn, Fe)S have equilibrated to
low temperatures (Skinner and Luce, 1971), andLarimer and Buseck (I974) deduced equilibration
at 91o-11 to K (possible systematic error of 15o K)

M I N E R A L O G Y OF THE PLANETS

65

for the Types I and II chondrites at Ps2 C.lO-S- has albite (Ab92AngOr4) and metal (inclusions
10 - 1 2 bar and Po2 C ' I 0 - 2 8 - I O - 3 7 bar.
with 5o ~ Ni set in matrix with 4.5 ~o). Enstatites
Although there are general similarities among are ordered orthoenstatite in unbrecciated Shallothe enstatite achondrites, detailed mineralogical water, but are sometimes disordered in brecciateddata indicate that sub-types will be needed.
ones, probably because of impact deformation
Olsen et al. (1977) distinguished the Happy (Pollack, I966; Reid and Cohen, 1967).
Canyon specimen as a separate type with a crystal
Systematic study showed a negative correlation
cumulate texture attributed to melting of EII between Si in metal and Fs in enstatite for somechondrite at a slightly higher oxidation state (to enstatite chondrites and achondrites (Wasson and,explain absence of Si in kamacite; i ~o Ti and 5 ~o Wai, 197o; Wai and Knowles, 1972 ). Metal parCr in troilite; enstatite Fso.4) in a body big enough ticles occur inside enstatite crystals, and the grain
to give congruent melting of enstatite. Perhaps the size of chondrites increases as the chondrule condiopside (EnsE.3Fso.9Wo46.8) exsolved out of a tent decreases. Wasson and Wai (197o) proposed
solid solution of an enstatite polymorph and the that a planetesimal accreted enstatite chondrules
pyroxenes equilibrated to c.II5O K based on the from a cold solar nebula with mechanical sepaLindsley and Dixon (I976) calibration. Congruent ration of metal and silicate, and that external
melting of enstatite requires a pressure greater than heating yielded achondrites on the outside, but notc.2 kb (Boyd et al., 1964). Perhaps the pressure was by igneous activity.
high enough to yield ortho-enstatite instead of
Based on the textural similarities of enstatite
proto-enstatite. The ptagioclase (An26) is more chondrites and ordinary chondrites, I tentatively
calcic than that in EII chondrites (An13_17), and is prefer to generate the former from a debris cloudsporadically abundant (5-IO ~o).
formed by planetesimal collision, but await
The Cumberland Falls meteorite has angular detailed study to test this idea. The pyroxene
blocks of black chondrite set in an enstatite-rich relationships may favour derivation of the achondbreccia indistinguishable from other brecciated rites at considerable depth, and I prefer to consider
enstatite achondrites (Mason, I962 ).
them very tentatively as brecciated cumulates.
Unlike the silica-bearing enstatite chondrites the Probably the present data are most easily interprincipal group of enstatite achondrites contains preted in terms of several planetesimals, butseveral per cent of forsterite in association with detailed mineralogical and chemical studies may
dominant enstatite and minor plagioclase, kama- allow coordination of some sub-types into a single
cite, and essentially the same group of accessory chemically zoned planetesimal.
minerals as in the enstatite chondrites. The unbrecBulk chemical data were interpreted in terms ofciated Shallowater meteorite may have a cumulate nebular condensation, metal-silicate-sulphide
texture (Foshag, 194o ) with equant to spherical mechanical fractionation, and migration of mobile
grains of forsterite (5 ~ ) and irregular grains of elements (Baedeker and Wasson, 1975; Binz et al.,oligoclase (2 ~o) poikilitically enclosed in enstatite T974; Ikramuddin et al., 1976 ) but analyses oflaths (81 ~o). Irregular masses of Ni, Fe (9 ~o) up to 3 separated mineral fractions are needed to test the
mm across show Neumann lines and pearlitic models.
texture, and are sometimes associated with troilite
Melting of a hypothetical planetesimal might(1.5 %).
have produced an E-achondrite mantle and an
All the other E-achondrites are monomict brec- iron-rich core, thereby explaining the low Fe and S
cias with coarse fragments of enstatite set in a content of E-achondrites. Only three iron meteorgroundmass dominated by enstatite. The Norton ites have Si in the metal (Tucson, o.8 w t ~o; Horse
County achondrite (Beck and LaPaz, I95I ; Keil Creek, 2. 5 ; Nedagolla, o. 14 wt ~o; Wai and Wasson,and Fredriksson, i963) has metal inclusions 5 cm 1969, I97O). Troilite is probably absent in Horseacross composed of Ni-poor kamacite (Ni 3.6 Creek, Tucson, and Nedagolla (Buchwald, I975),wt~), and disseminated troilite (Ti o.4-4 wt~), but Tucson contains brezinaite (Cr3S4) and abunalong with copper, alabandine, daubr6elite dant silicate inclusions (Bunch and Fuchs, I969).
(FeCr2S4), and graphite. The Mayo Belwa ach- Although these inclusions Contain abundant
ondrite (Graham et al., 1977a ) has highly shocked forsterite (FeO o.23; CaO O.ll Wt~o) and minor
enstatite (Woo.6; FeO o.o2 wt ~o), forsterite, diop- orthopyroxene (En99Fso.4Woo.6; A1203< o.o3
side (En55FSlWO~4; hence c.I35o K from Wt~o), which might be matched with enstatite
Lindsley-Dixon calibration), albite (AbaaAn9Ora), meteorites, the rare diopside (En54.8Fso.sWo44.7)
Si-bearing metals (with Ni 2.5, 7.6, and 2I Wt~o), contains 7.3 wt ~o A12Oa and the feldspar is pure
Cr,Ti-troilite, daubr6elite, oldhamite, and fluor- anorthite. The Tucson iron has a texture consistentamphibole needles projecting into 1 cm vugs. The with rapid cooling of taenite (Buchwald, I975).
Khor Temiki achondrite (Hey and Easton, I967) The Nedagolla iron was remelted (Miyake and66

J . V . SMITH

Goldstein, i974) , and both it and the Tucson iron
were cooled very rapidly (Nedagolla I K/rain;
Tucson I K/moo yr).
The Bencubbin metal-silicate breccia contains
clasts related to carbonaceous (c.CM-CO) and LL
chondrites (Kallemeyn and Wasson, ~977) plus
clinoenstatite and rare olivine crystals set in amatrix of opaque silicates (Newsom and Drake,, I977). Abundant metal clasts are rich in Cr and S,
and preliminary analyses indicated Si substitution
in some.
Obviously much study is needed to explain thechemistry, origin, and development of the enstatite meteorites and reduced irons.
Forsterite-bearing meteorites and silicate inclusions in irons. The iron contents of olivine andpyroxene are a guide to the extent of oxidation,
and, in the preceding sections, the olivines haveranged as follows: pure forsterite crystals in CII
meteorites (perhaps early condensate from thenebula) to unequilibrated Fe-bearing olivines inother C-meteorites; unequilibrated Fe-bearing oilvines in type 3 ordinary chondrites to equilibrated
ones in the higher grades (H,Fals; L,Faz3;
LL, Fa28), and almost pure forsterite in E-achondrites.
Graham et al. (I977b) described four meteorites
with substantial Mg-rich olivine and orthopyroxene, and similar compositions of troilite andmetal. The Kakangari chondrite was briefly mentioned earlier as a member of the tentative CKgroup of carbonaceous chondrites. The remaining
three meteorites lack chondrules, and have chemical and mineralogical analogies with silicate inclusions in Type I iron meteorites (Bild, I977; Davis etal., I977). The Pontlyfni meteorite (Graham et al.)has abundant sulphides (21 ~o troilite, daubrrelite)
and metal (25 ~o kamacite, taenite) associated withsilicates (forsterite, enstatite, diopside, oligoclase) ina fine-grained mosaic. The Mount Morris (Wis)and Winona stones are badly altered, but containminerals of similar composition, as well as graphite
and chromite (refs. and data in Bild, 19771.
Silicate inclusions occur in half of the group-IAB
irons (Buchwald, I975), and probably result fromtrapping in pools of liquid metal. Most othergroups of irons do not contain silicates, but a few
anomalous and liE and IlICD irons do. Bunch etal. (I970) described three groups and three single
irons, and further data are summarized in Bild
(1977). The Odessa (IA) type of silicate inclusions
has > 5 ~o graphite and troilite, rounded to irregularshape, granoblastic texture, olivine Fa~4, enstatite
En91_92WOl.2__l.9, diopside En51_saFs3Wo4~46(hence equilibrated to sub-solidus temperaturesc. ~3oo K), plagioclase Ang-i4Or~4, zincian chromite of distinctive composition (Bild, t977), troilitewith o.2-o.6~o Cr, alabandine, sphalerite with
22-3I wt~o Fe, Cl-apatite, merrillite, kamacite,taenite, and copper. [Whitlockite was t h e nameused in most papers up to I978 for the calcium
phosphate mineral in meteorites. Because of small
structural differences from terrestrial whitlockite,
the name merrillite has been resuscitated by the
I.M.A. Commission on New Minerals and Mineral
Names for the lunar and meteoritic mineral. Merrillite has been used here for all meteoritic occurrences, but definite proof is needed especially for
CI meteorites.] The Copiapo (IA) type has< 5 ~ graphite and troilite, angular shape, diopside En52_55Fs2~,W042~4,
enstatite En90_95
Fs4__9Wol_2, forsterite Fal_7, plagioclase An11_22
Or2~ , Cr-spinel, and Cr-troilite. Scott and Wasson (~975) did not accept the distinction between
Odessa and Copiapo types. The alkali-rich Weekeroo Station (IIE) type has no graphite, variable
texture, drop-like inclusions, augite En46Fs~s
Wo36, bronzite En75Fs22Wo3, no olivine, Crspinel, Cr-free troilite, phosphate, and K-richfeldspars (Bunch and Olsen, I968). The shocked
Kodaikanal member (Bence and Burnett, ~969)
contains high-albite solid solution (mb61_86
Or11_39with o - I o ~ excess SiO2), glass Ab42Or44
Qz~4, augite En52FsloWo3s, orthopyroxene Ensl
FsavWo2, chromite, Ca-phosphate, and kamacitetaenite Widmanst/itten intergrowth indicating
rapid cooling of I02-Io 3 K/Io 6 yr. Enon (Anom)with Fe-rich olivine and orthopyroxene, and oligoclase Anx 5Or6, resembles mesosiderites. The Steinbach stony-iron has metal of IVA type, and thesilicate portion (c.50 ~o) contains tridymite, orthopyroxene En85(CaO o.23 wt ~o) and clinopyroxene
inverted from protopyroxene (Enss; CaO 0.I7),
chromite, troilite, and schreibersite. The original
assemblage should have equilibrated at 147o K,fo~
IO-~2 bar, and total pressure < 2 kbar; rapid
cooling at high temperature is indicated by theclinopyroxene texture, and slow cooling (5-2oK/io 6 yr) at low temperature by the diffusion
profile of the Widmanst/itten pattern and theMg, Fe ordering in the orthopyroxene. Kendall
County (Anom) has graphite with minor Fe-freeolivine and pyroxene, and variable plagioclase
Absv-Ab51. Netscha~vo (IIE-Anom) has recrystallized relict chondrules, diopside Ena9Fs6Wo45 ,enstatite Ens4FSl4WOl.o4, oligoclase An14Or4,
chromite, merrillite, chlorapatite, metal, troilite andschreibersite (Olsen and Jarosewich, I97i ).
Bunch et al. questioned why the pyroxene
thermometer gave temperatures (cA300-~500 K)
much higher than the ones required to give the
kamacite-taenite assemblage (< 7oo K?), and suggested trapping of metamorphosed silicate inyounger Ni-Fe melt. However, this problem also

M I N E R A L O G Y OF THE PLANETS
occurs in other meteorite groups and may result
from easier solid-state diffusion in metal than insilicate.
The Winona meteorite contains zincian chromite

67

and silicates of similar compositions to those in
the Odessa-type irons. The following table showsother comparisons between forsterite-bearing
meteorites:

Kakangari
Mt. Morrison Winona
Pontlyfni
olivine (Fa at ~o)3-9
35
o.8-I.3
orthopyroxene (at ~o)
Fs3_14Wo2
Fs4Wo2FsTWo2
Fsl_3Wo2
diopside (atTo)
not found
FslWo46-FslWo46
feldspar ( a t e )-An23Or3
An16Or4
An12Or3
troilite (Cr wt ~o)
0.3
0.4
-0.8
kamacite (Ni Wt~o)
6.o
6.o
6 to
6.I
taenite (Ni wt ~o)
26-32
35-44
14
23-28
total Fe (wt ~o)
23
20
-34
The Mundrabilla iron (IA-Anom) is remarkable Ge (Scott and Wasson, 1975). All members of afor the high S content resulting from c.35 volvo G a - G e group are believed to arise from the same
troilite nodules (Ramdohr, I976; Bild, I977), some planetesimal, while the range of textures (if any)
containing angular silicate inclusions (olivine Fa3, gives some clue to the physical locations of theorthopyroxenc Fs6.sWo2, clinopyroxene Fs3Wo45, irons in the parent body. Some irons were severely
plagioclase An 11Or4). The troilite contains o.5 wt ~o reheated. That many have been shocked (Jain et al.,Cr and o. I/oo, Cu (Weinke, I977) and other sul- 197z; Wasilewski, J 976) is not surprising since irons
phides are daubrrelite, alabandine, and sphalerite. should tend to be deep-seated in differentiatedGraphite occurs mainly at grain boundaries, often planetesimals.
Principal textural classes are hexahedrites (H)associated with schreibersite. The Watervillemeteorite (Scott and Wasson, t976) is minera- [large single crystals of kamacite ( = ct = bodylogically similar, but finer-grained and lower in Ir centred iron); 5.5 wt ~ Ni; G a - G e group IIA; somerecrystallized to fine grain size upon reheating],(o.3o ppm vs. o.87).
The Campo del Cielo (IA) iron has mineralogical octahedrites [kamacite-taenite intergrowths atand chemical properties indicative of heating to a tached on the octahedral planes of the original
higher temperature than other IA irons with silicate high-temperature single-phase taenite ( = "; --- facecentred iron); the Widmanst/itten structure varies
inclusions (Wlotzka and Jarosewich, 1977).
Unquestionably the discovery of silicate inclu- in coarseness from coarsest octahedrite (Ogg) via
sions in irons depends somewhat on chance. Tem- Og(coarse), Ore(medium), Of(fine), and Off(finest)
porarily it seems best to assume that the IAB irons to plessitic octahedrite (Opl)], ataxites (D) [heteroare characteristically associated with silicates geneous, fine-grained; Ni-rich ones are kamacitewhose Mg/Fe ratios indicate aggregation into a taenite intergrowths] and anomalous irons [often
planetesimal from material of oxidation state inter- fine-grained, or with abundant sulphide or silicate,
mediate between those for ordinary chondrites and or both]. The over-all features can be explained byenstatite chondrites. Perhaps the IAB irons, the Mt. the binary relations between low-temperature
Morrison, Winona and Pontlyfni stones, and the kamacite and high-temperature taenite in the FeMundrabilla troilite-iron, are samples of a modally Ni phase diagram (reviewed by Goldstein andheterogeneous planetesimal composed of collision Short, I967), but many complexities result fromdebris that became hot enough for partial melting nucleation features related to minor substituents
but not core separation. Other bodies are needed to (especially, P, C, and S; e.g. Clarke and Goldstein,accommodate the Weekeroo Station (liE) irons i978 ). Diffusion profiles lead to estimates of cooling rate, which are based on the key assumption ofand anomalous irons not in the IAB group.
Irons. Iron meteorites are valuable indicators of monotonic cooling.
Important chemical features of the twelve major
the accretion processes and metal-sulphide-silicate
differentiation of at least a dozen and perhaps sixty G a - G e groups are listed in Table V taken mainly
planetesimals, or regions of planetesimals, for most from Buchwald 0975; Table 27), but extended withof which there are no obvious silicate-rich meteor- data from Scott and Wasson (I976, Table 2).
It is quite impossible to do justice here to theites. This results from the augmentation of thirtyone fallen irons by c.53o finds, and the lesser magnificent compilation of mineralogical and texfraction of finds of stony meteorites. The Heroic tural data in Buchwald (I975) together with theAge produced a textural classification, which corre- detailed metallographic and geochemical studies ofsponds roughly with bulk Ni, and a burgeoning many workers. From the viewpoint of planetarychemical classification based principally on Ga and mineralogy emphasis is placed here on features

68

J. V. S M I T H
TABLE V. Summary of iron meteorites

Ga-Ge
group

Approx.
numberTextural
class

Cooling*
rate

Ni
wt %

Ga
ppm

Ge
ppm

Ir
ppm

IA
IB?
IC~:
IIA
IIB
IIC
IID
IIE
IIIA
IIIB
IIIC
IIID
IIIE
IIIF
IVA
IVB
Anom

82
8
IO
43
14
7
1I
12
i 17
4o
6
5
7
5
39
II
92

Om-OggD-Om
diverse
H
Ogg
Opl
Of-Om
variable
Om
Om
Of
Off-D
Og
-Of
D
All

2 3
-1-1oooo
2; Io-lOO-5OO

6-9
9-25
6- 7
5.3 5.7
5-7
9-I2
IO-II
7-1o
7-9
9-11
I I- 13
16- 23
8-9
7-8
7-9
16 18
5-60

55-100
11-55
42-54
57-62
45-60
35-40
70 8 5
21-28
17-22
i6 22
11-26
1-5
17-19
6.3-7.2
1.6-2.4
o.17 0 . 2 7
0.05-87

I9O-52o25 19o85-2500.6-5.5
0.3-2.0
0.o7-92 60
O.Ol-O.5
4-I1
3 20
0-5-8
o. 15-20
o.0i-o.15
0.07-0.55
O.01-0.07
0.05-0.6
1.3-7. 9
o.I-4
4 36
0.007-54

--

2
2-1o
---

-5-2007-2~176
5-200?
--

I70-I85

lO5-185
85 11580-1OO
62 75
32 -46
28-46
8-35
I-4
34-36
0.7 I.Io.o9-o.14
o.03-o.08
0.005-2000

Data mainly from Buchwald (1975, Table 27); see also Wasson (1974, Table 11--7), Scott and Wasson (i975, Table 2;1976) and Scott (I978c).
* Quoted as K/lO6 yr assuming monotonic cooling through 773 K; data from Kelly and Larimer (i 977) and Randich
and Goldstein (1978).
I See Buchwald (1975) for details of I-Anomalous irons listed here in group IB.
Scott (1977d).
w Disputed by J. I. Goldstein and J. T. Wasson groups.
resulting from accretion and from metal-sulphidesilicate differentiation.
Scott (I972) interpreted the trend between Ni
and various trace and minor elements of a G a - G egroup in terms of igneous fractionation within asingle parent body while the distinctions between
G a - G e groups were attributed to primary fractionation in the solar nebula. O f course, this binary
distinction c a n b e blurred if some planetesimals
accreted inhomogeneously. Furthermore the model of progressive equilibrium condensation o f thesolar nebula may be far too simple. For the latest
compilation o f trace elements, Scott 0978a) concluded that the non-members o f the twelve major
G a - G e groups are sparse samples of some fifty
additional planetesimals.
Kelly and Larimer (I977) set up a detailed
chemical model. At IO-5 bar the first condensate
contains Fe-rich metal with I 9 + 3 wt 9/0 Ni. It isenriched in all metals more refractory than Ni (viz.
Os Re M o Ir Ru Pt Rh), and strongly depleted inelements (Cu Zn Au G a Ge) more volatile than Ni(Wasson and Wai, 1976; Scott, 1978b). Equilibrium
cooling results in dilution of the Ni and refractoryelements with condensing iron to give c.6 wt % Ni.
The refractory elements are also diluted, while thevolatile elements are augmented up to the cosmic

ratio. Cobalt and Pd condense simultaneously withNi, and the C o / N i and P d / N i ratios remain constant. All the above elements, except Cu and M o athigh temperature, are more noble than Fe, andoxidation of Fe (with incorporation into silicate)
results in an increase in Ni content, while ratios ofnoble elements stay constant. Sulphidization of theFe will also increase the Ni content and not change
the ratios of non-chalcophilic elements; however,the sulphide could remove chalcophilic elements(e.g. Cu, Zn). Melting and gravitational segregation
in a planetesimal is very complex, but partial
melting can yield an Fe,FeS eutectic melt, which isgravitationally segregated from immiscible silicate
liquid or solid silicate residue to yield a core.
Alternatively, segregation is incomplete leaving'raisins' of Fe,FeS in a silicate host, or melting does
not occur, and there is merely solid-state partitioning of elements. Crystallization of the Fe,FeS
melt is controlled by the partition of elements
between solid and liquid phases. Chalcophilic elements should concentrate into the sulphide melt.
Elements that melt at high temperature should
favour solid Fe-rich metal over Fe-rich liquid.
Controlled experiments on these phenomena are
only just beginning by M. J. Drake, J. I. Goldstein,and co-workers, and present interpretations must

M I N E R A L O G Y OF THE PLANETS
be based largely on the element correlations observed by Scott for the various G a - G e groups.
All the above processes are complicated by theredox state, and by the nucleation phenomena
associated with minor amounts of phosphides
and carbides.
The four elements in Table V are sufficient to giveclues to the origin of the major groups of ironmeteorites, though readers are referred to many
details in published papers.
Group IVB is enriched in refractories (Ir 4-36ppm) and depleted in volatiles (Ge o.o3-o.o8 ppm),
and has high Ni 0 6 - I 8 wt~). Kelly and Larimer
suggested accretion at I27o K shortly after metalbegan to condense from the nebula. Minor metalsilicate fractionation may have occurred. The finetexture precludes accurate measurement of thecooling rate, but rapid cooling is indicated. GroupI I I D also has high Ni 06-23) but the Ir content(o.oi-o.o7) is much too low for a simple refractory condensate, and strong solid-liquid fractionation in a planetesima/is indicated.
Group IVA is strongly depleted in volatileelements but has low Ni (7-9) and Ir (o. I-4). Traceelements correlate with Ni as expected for fractional crystallization.
The abundant IIIAB irons have element correlations indicative of fractional crystallization ina single core, but the range of cooling rates z to
IO K / i o 6 yr may indicate more than one metalreservoir. These irons resemble ordinary chondrites
in having five-fold depletion of volatile elements(e.g. Ga, Ge), a similar redox state, and exposure
ages indicating a major collision at - 6 . 5 x Io 8 yr.
Detailed study of the huge Cape York iron revealed
many trace minerals (Kracher et at., 1977): four Nabearing phosphates (including buchwaldite
NaCaPO4), six sulphides (daubr6elite, chalcopyrrhotine, djerfisherite, another K-rich sulphide,sphalerite, Fe-alabandine), chromite, copper, carlsbergite (CrN), and a silica mineral. The main groupof pallasites (next section) is chemically related tothe high-Ni IIIAB irons (Scott, I977a , c).
The abundant IIAB irons have low Ni and highG a and Ge indicative of low-temperature accretion, and strong Ir variation indicative of strong
crystal-liquid differentiation in a single core. Six
IIA hexahedrites have cooling rates of 2 K / i o 6 yr
whereas one has IO K / I o 6 yr (Randich and Goldstein, I978). The low Ni content of 5.3-5.7~ forhexahedrites indicates that all the Fe and Ni were in
the metal phase, and none was in a silicate phase.
However, hexahedrites are not as reduced as enstatire meteorites because they do not contain Si in themetal.
Group IAB irons have trace element trends
interpreted by Kelly and Larimer as the result of

69

seepage of metal during partial melting but without
aggregation. These irons are characterized by silicate inclusions described in the preceding section.
The angularity of the silicate inclusions suggests
derivation from a metal-silicate agglomerate.
Gravitational segregation of metal and silicate
occurred only locally. S01id-state metamorphism
yielded a silicate-equilibrated mosaic with pyroxene grains established at c.I3oo K. The kamacite and taenite intergrowth results f r o m solid-state
annealing at c.6oo K of a taenite solid solution
established at > Iooo K.
Group IC irons (Scott, I977d) have low Ni andhigh Ge and Ga, together with an I r - N i correlation
indicative of metal-liquid fractionation. Cohenite
and chromite are abundant and the cylindricalsulphide inclusions are free of silicate and graphite.
Schreibersite and minor carlsbergite occur. The
wide range of cooling rate ( I - I o 4 K / I o 6 yr) was
attributed to burial at different depths of impact
debris from a body undergoing metal-liquid fracfionation.

It is convenient here to mention briefly theevidence on the pressure at which meteorites
crystallized. The early conclusion by Anders (t 964)
that the maximum static pressure was not abovec.3 kb still stands for at least the great majorityo f meteorites. This conclusion rests on the lack of
inversion of tridymite to quartz, the evidence for
protopyroxene, the attribution of diamond toshock rather than static pressure, and the reinterpretation of early speculations that high pressure
was needed for cohenite (Brett, x967) and kamacite-taenite intergrowths (Goldstein and Doan,I972). For meteoritic silicates there is no need to
invoke static pressures greater than a few kilobars,
except just possibly for pallasites (next section). For
irons, the Fe content of sphalerite, when in equilibrium with troilite, provides a barometer if thetemperature can be estimated. Schwarcz et al.
(I975) estimated pressures of 1.7-3.I kb for the
Bougou, Gladstone, Odessa, and Sardis irons ofType I. The pressure increases as Ge decreases,suggesting a chemically zoned planetesimal ofradius up to c. 18o km if made only of iron and c.35o
km if silicate-rich. Further studies of sphaleritebearing irons have not yet appeared, except for an
estimate of IO-I 5 kb for the Cape York IIIA iron
(Kracher et al., I97 % and c.1 kb for the sphaleritewith 28 wt ~ Fe in Mundrabilla (Ramdohr, t976).
The importance of estimating the cooling ratefrom kamacite-taenite intergrowths was recognized by Wood (t964) , and later work has demonstrated the complications caused by supercoolingand heterogeneous nucleation on inclusions (reviews by Goldstein and Short, x967; Goldstein and
Axon, I973). Controversy continues over details,

70

J . V . SMITH

but Widmanst/itten intergrowths must have exsolved from a supercooled parent taenite solid
solution at a temperature below c.9oo-IOOO K that
depends on the bulk Ni content. Furthermore, thecooling histories of iron meteorites must varyconsiderably, and the possibility of complex cooling histories, such as annealing in a breccia after
disruption of a planetesimal must be considered.
Utilization of phosphide inclusions in hexahedrites(Randich and Goldstein, 1978) points the way tofurther estimates of cooling rate using exsolution ofminor-elements, while new analytical techniques
may permit study of fine-grained plessites (Lin etal., I977).
The occurrence of S, C, P, and N in ironmeteorites is particularly relevant to planetarydifferentiation. Troilite is the most abundant sulphide mineral. Complete separation of an Fe,FeS
eutectic might be expected in differentiatedplanetesimals with subsequent crystallization into
metal and troilite, which might become mechanically segregated in a gravitational field. Meteorites
composed entirely of troilite are not known,
perhaps because fragmentation of a planetesimalcore would shock-melt or vapourize a troilite
region. Most iron meteorites carry little troilite( < 2 ~), but a few contain a lot, of which theMundrabilla (preceding section) and Soroti irons
are important examples (Buchwald, I975). It would
be very useful to attempt to estimate the metalsulphide ratio from the relative content of chalcophilic to siderophilic elements in iron meteorites,
but different condensation properties from thesolar nebula cause problems. Taken at face value
the planimetric data of Buchwald (1975, Table 3o)indicate that the S content of most iron meteorites
is much less than expected for the cosmic Fe/S
ratio, but it is not certain whether this results fromlower condensation of volatile S than less-volatile
iron, or sampling effects. There is no obvious
correlation between S content and assignment ofiron meteorites to either a core, raisins, or incipient
melt regions, except for the high S content of IAnom irons attributed to non-segregated melts.
Troilite is absent in the reduced irons Tucson and
Nedagolla, but brezinaite occurs in the former anddaubrrelite in the latter.
Carbon is sparse in irons except for group I witho.2-2 wt 9/0 graphite, cohenite, and haxonite, andgroups IC, IIICD, and IIIE with somewhat lower
abundances of carbides (Scott and Wasson, 1975,
Table 3). Graphite is commonly associated withtroilite, while cohenite occurs in metal, especially in
IC irons. In spite of this low abundance in irons
simple analogy suggests that an iron-rich core may
be the major store of C in terrestrial planets, rather
than the atmosphere and crust (Smith, 1977).

Nitrogen has a very low abundance in irons (2I3O ppm; Gibson and Moore, 197i), but simple
analogy again suggests that a planetary core could
be the major reservoir rather than the atmosphereof terrestrial planets. Carlsbergite, CrN, occurs
particularly in low-Ni irons of groups IIIA, IIA,
and I (Buchwald, 1975).
Phosphorus is important in planetary differentiation as an indicator of oxidation state.
Schreibersite (Fe,Ni,Co)3P, and its textural variety
rhabdite, are almost ubiquitous in iron meteorites(Doan and Goldstein, 1969; Scott and Wasson,
1975, Table 3) and arise either from P-rich liquid or
exsolution in the metal. This proves that the irons
developed in an environment sufficiently reducingthat the P was not dominantly in the oxidized form
(as in merrillite). Rare phosphates, some associatedwith phosphide, occur in IIIAB irons (Buchwald,1975) and in the Verkhne Dnieprovsk and Barranca Blanca IIE (Olsen and Fredriksson, I966;
Bild, 1974) and Dayton IIID irons (Fuchs et al.,1967); four of the phosphates contain sodium(Kracher et al., 1977). The implications for theoxidation state await experimental synthesis, andcalculation of possible oxygen diffusion.
The highly reduced state of the irons containing
Si-bearing metal and Ti-bearing troilite was described earlier. Chromium is probably a common
minor constituent of troilite from iron meteorites,
whereas Ti probably occurs only in the very
reduced ones. Systematic study of these and otherelements in troilite and other sulphides from irons
might provide a guide to the redox state, but thecomplex sulphide mineralogy must be taken into
account (e.g. E1 Goresy, 1965, I967) when interpreting distribution coefficients.
From this brief and highly selective survey it isabundantly clear that iron meteorites derive from avariety of nebular condensates that have been
modified in many ways, including differentiation in
many planetesimals. Further study, especially of
anomalous irons, should increase the evidence fordiverse origins.
Pallasites. These stony-iron meteorites consist
essentially of olivine and metal, whose complex
textures require prolonged solid-state annealing.
Those with c.65 vol ~o olivine were interpreted by
Buseck (I977) as cumulates of olivine crystals
whose interstices were filled gently by liquid metal,
but Scott (I977b) argued for violent mixing ofmolten metal with brecciated olivine. Those with> 65 vol ~o olivine were interpreted as adcumulates
by Buseck (i977). A 'mean' pallasite has 65 volvo
olivine, 31 ~o metal, 2. 3 ~o troilite, 1.2 ~o schreibersite, o.4 ~o chromite, and o.z ~ phosphate giving abulk composition: P205 < o.I, SiO2 19.7, Cr203
o.2, FeO 6. i, MgO 21.o, Fe 44-4, Ni 5.7, Co o.2, S o.2,

M I N E R A L O G Y OF THE PLANETS
P o.3 wt ~. Most pallasites contain metal withsimilar Ni, Ga, Ge, Cr, W, As, and Au to high-Ni
IIIAB irons (Scott, I977a, c), and the chemical andtextural properties of the latter would be consistent
with derivation from the core of a planetesimal ofwhich the pallasites represented the core-mantle
interface. However, the apparent cooling rates
inferred from the Widmanstfitten intergrowth are 2
to IO K/io 6 yr for the irons and about I K/Io 6 yr
for the pallasites (Buseck and Goldstein, 1969).
Perhaps this can be resolved by appeal to different
annealing conditions after disruption of theoriginal parent planetesimal, but mineralogicalevidence might be difficult to interpret. Three
pallasites (including Eagle Station) have high
Ge/Ga, and may be akin to some anomalous irons.
Six pallasites are unique.
Buseck (1977) emphasized cumulate texture,
angular shape, kink-banding of olivine, troilite
eutectic with kamacite, and deformed plessite.
Scott (i977b) emphasized disintegration of olivine
masses into smaller angular fragments as the resultof intrusion by molten metal. Adjacent olivine
grains inside the fragments have triple contacts
typical of solid-state annealing while olivines incontact with metal may have many small faces
defining a nearly rounded interface. Scott suggested that the gravitational separation of olivine
and metal was probably prevented by incipientsolidification. Wood (I978) concluded that because olivine deforms easily olivine-metal cumulates would squeeze to pure dunite and metal inbodies of radius > IO km. In order to obtain theslow cooling indicated by the Widmanstgtten texture it would be necessary to aggregate small solid
bodies into > IOOkm bodies before final cooling of
pallasite through 8o0 K. Probably it will be necessary to consider at least the following processes:
melting, cumulation, deformation and solid-state
annealing, violent disruption, further solid-state
annealing, and perhaps final fragmentation. Cooling was probably not monotonic, and estimationof radii of parent bodies may be difficult or
impossible.
The olivine and metal have Fa10.s-13 (atomic)
and Ni8-I2 wt ~ (main group), and Fa19_20 and Ni
I4-I5 (Eagle Station trio). Olivine grains containtubes I-5/~m wide, and in tiny regions contain 4-5
wt ~o PzO 5 (Buseck, 1977). Rare symplectites occur
with pyroxene (Fs~ z or Fsl 7). Cobalt is depleted 7ootimes in the olivine compared to metal, whereas Ni
is depleted c.7ooo times (Turekian et al., I977). Thisis the prime basis of the argument that the Earth's
upper mantle (not necessarily the lower mantle) did
not equilibrate with liquid metal. Kolomeitseva(I975) interpreted Fe, Mn, Ni, Cr, and Co of
pallasitic olivine as the result of metal-silicate

7I

equilibration at I4oo-I5oo K and 20-70 kb, but
this cannot be accepted until partition coefficients
have been measured on synthesized phases.
Phosphorus occurs as widespread schreibersite
and trace phosphates (merrillite with minor Na andK, stanfieldite, and farringtonite; Buseck andHoldsworth, 1977). Trace minerals include low-Ti
chromite, sphalerite, pentlandite, copper, andgraphite.
Chassignites and nakhlites. In view of all theevidence for differentiation, it is surprising that
there are so few olivine- and pyroxene-richmaterials. Perhaps olivine and pyroxene tend tofragment upon impact unless bound together byother minerals.
The Chassigny meteorite (Mason et al., 1976;
Floran et al., I977a ) is texturally a shocked cumulate of high-Fe olivine (Fa32; high CaO o. 16 wt ~o)with minor low- and high-Ca pyroxenes, plagioclase (Ca 4.63 N a 2 0 8. I5 K 2 0 ~.47 wt ~o), sanidine
and chromite, and traces of chlorapatite, troilite,hydroxy-fluor-kaersutite (in melt inclusions inolivine), pentlandite, ilmenite, rutile, and baddeleyite. Bulk chemical data (Boynton et al., 1976) fit
with an olivine-rich cumulate from a chondritic
magma.
A second chassignite, the recently discovered
Brachina meteorite (Johnson et al., I977), contains
dominant olivine (Fa33; CaO o.25 wt~o) in a
subhedral granular texture along with minor augite(En46Fs14Wo40) and plagioclase (CaO 4.71 N a 2 0
8.92 K 2 0 o.26 wt~o) and accessory chromite,
troilite, pentlandite, and inferred phosphate.
Olivine and chromite grains project into voids, andthe meteorite is porous and unshocked.
The Nakhla and Lafayette meteorites are identical pyroxenites (Bunch and Reid, 1975; Boctor etal., 1976) with cumulus augite En38Fs23Wo39,
minor olivine (Fa65_68; CaO 0.45), interstitialoligoclase-sanidine (An34Or4-OrvsAn3), Cr-Timagnetite with exsolved ilmenite, pyrite withmarcasite lamellae, troilite, chalcopyrite, and F-C1apatite.
The Fe-rich nature of the olivines is notable in all
four meteorites, and it may prove possible tointerpret the nakhlites as later differentiates thanthe chassignite in a planetesimal of Fe-rich peridotitic composition, perhaps chemically related to
low-Fe chondrites. Laul et al. (1972) found thevolatile elements of nakhlite to resemble those ofocean-ridge basalts, and Gale et al. (1975) reaffirm a
Rb-Sr isochron that indicates igneous crystallization at - 1.24 x I09 yr.
Angrite. The unique Angra dos Reis achondrite(Hutchison, ~972; consortium of papers, see Prinz
et al., I977b) has 93~0 fassaite (Wo54En34Fs12;
TiO 2 2.2 A120 3 io Wt~o) occurring in a texture

72

J.V.

SMITH

interpreted as oikocrysts set in a lineated groundmass. Triple j unctions were interpreted as the resultof metamorphism following igneous cumulation.
Rare olivine (Fos3,CaO 1.3 Wt~o) occurs with Mgkirschsteinite, hercynitic spinel (requiring Fe3+),
merrillite, Ni-iron, troilite, plagioclase (An13Oro.7),
celsian (Cn9oAnaAb2), baddeleyite, and Ti-magnetite. High pressure is not required to explain the A1
content of the pyroxene. A parent body of unusualcomposition is required for this meteorite, perhaps
very rich in high-temperature condensate (i.e. like
Ca-Al-rich inclusions in CIII meteorites).
Ureilites and lodranite. The Lodran meteorite(Bild and Wasson, 1976; Prinz et al., I978 ) contains
three-eighths olivine (inversely zoned to grainboundaries and veins from Fa12 to Fa6) and threeeighths unzoned orthopyroxene (Ena4Fs13Wo 3with exsolved augite EnsoFsvWo43), interpreted asa cumulate harzburgite surrounded by one-quarter
metal (mainly kamacite; minor taenite) and troilite.
Minor phases are schreibersite (which yields a
cooling rate of io K / I o 6 yr), chromite and 'granitic'
inclusions. The olivine zoning indicates reduction
after crystallization, and various chemical features(including oxygen isotopes) indicate a different
parent body than ureilites in spite of similar
mineralogy. Fukuoka et al. (I978) interpreted bulk
chemistry in terms of partial melting of a hybrid
between H and E chondrites.
The seven ureilites have remarkable properties,
which involve a complex history and controversial
interpretations. Berkley et al. (1976, 1978 ) describe
them as a pile of igneous cumulates with graphite(probably primary), which caused reduction ofolivine (as for the Lodran meteorite). Strong shock
produced diamond and lonsdaleite. The ureilites
were classified as:
wt ~o F e O /(FeO + MgO) wt ~o Cgroup i(top)

Dyalpur, Dingo
Pup Donga,
North Haig
'group' 2 Hayer6
group 3
Novo Urei,(bottom) Kenna, Goalpara

O.I8 0.22
0.27

3.6
2.I

0.25-0.32

1.5-2.2

Wasson et al. (1976) also interpreted the ureilites asa section of a planetesimal but from measurements
ofNi Zn Ga Ge Cd In Ir and Au placed five ureilites
in a different order from top to bottom: Goalpara(heavily shocked, low Ir and Ir/Ni), Hayer6, Dyalpur, Novo-Urei, and Kenna (moderately shocked,
high Ir, high Ir/Ni). External heating was supposed
to have caused partial melting with greater extraction of metal in Goalpara than in Kenna. A major
impact then injected C-rich and metallic material,
which was not primary as in the model ofBerkley etal., into hot silicates. Olivine zoning requires reduc-tion and rapid cooling. Diamond formed either
during injection or later. The primary metal was ahigh-Ir nebular condensate, partly lost by partial
melting, and the injected metal was low in Ir.
Boynton et al. 0976) from bulk chemistry interpreted the ureilites as a mixture of ultramafic
silicates with higher Co/Ni than metal, an indigenous metal phase remaining after partial melting,and carbon-rich material added after partial melting. The C-rich material is remarkable not only forthe diamonds and lonsdaleite, but also for beingrich in volatiles and primordial gases (Weber et al.,i976 ). CI-type material was ruled out by Binz et al.
(I975) , who suggested a late 'distillate' of mobile
elements. Higuchi et al. (1976) observed ten-foldenrichment of refractory Ir and Re over moderately
volatile Ni and Au, and paucity of volatiles except
for Ge, C, and noble gases. They interpreted
ureilites as the residue from partial melting of a
C3V-like body with incomplete separation ofliquid. Vein material was injected later. Carbon wasderived by surface catalytic reactions at c.5oo K.
Noble gases but few other volatiles were trapped.
Clayton et al. (I976) explained the oxygen isotopes
by mixing of e.o.2 ~o t 6 0 with anhydrous phases in
C2 and C3 meteorites, or similar material.
Returning to mineralogy and petrography
Berkley et al. (I976) provided detailed evidence inthe Kenna ureilite for cumulation of olivine (coreFazl,CaO o.42, Cr203 o.63 wt ~o; rim to Fal) andpigeonite (En73Fsl sWo9) in a lineated texture withtriple junctions. Secondary melt inclusions andveins in olivine and pigeonite contain clinopyroxene (Ensz_6oFs6_zsWo22~,2), K-feldspar,
andesine, chromite, and glass. Thin metal veins andblebs are rich in Co (mean 1.2 wt ~o) and have Ni 49 wt~o. The carbonaceous interstitial material
contains graphite, diamond, and lonsdaleite, andoften occurs as veins tapering into metalliferous
extensions, which contain some troilite and taenite.
Brecciation and extreme mosaicism of olivine
occur in Goalpara, Dingo Pup Donga, North Haig,
and Haver6 ureilites, but not in Kenna. Texturalfeatures in the Kenna silicates indicate shock in the
5o 25o kb range, which should be high enough toform the diamond and lonsdaleite.
There can be no doubt that melting has occurred,
and that rapid cooling and shock deformation
followed, but the other events are highly controversial. Probably the graphite is not a cumulate
phase because it would not be in equilibrium with
Fe-rich olivine. Introduction as a vapour phasewith grain-boundary reduction of olivine, perhaps
at low temperature over a long period, seems more
likely, especially as this could coincide withentrance of rare gases. Perhaps a shock event
tended to fracture the igneous cumulate along

M I N E R A L O G Y OF THE PLANETS
grain boundaries, thereby providing channels forthe vapour phase (note that eclogite nodules fromkimberlite typically are altered in the grain boundaries). The lineated texture may be difficult to
explain in a residue from partial melting unlesssome directed stress were available; igneous
cumulation seems more plausible. Perhaps theureilites result from six events: igneous processes ina parent body containing some 160-rich material;
disruption and incorporation as hot debris in
another parent body; vapour deposition into fractures, mainly along grain boundaries; further disruption with formation of diamond; rapid coolingof the shock debris; and transport to Earth.
Eucrites and shergottites. These two groups ofbasaltic achondrites are mostly monomict breccias
dominated by plagioclase and high-Fe pyroxenes,
but a few unbrecciated specimens occur. The earlysurvey by Duke and Silver (I967) outlined theevidence for magmatic crystallization followed byvarious degrees of impact and thermal metamorphism, and recent detailed studies are profiting by
experience on lunar rocks.
The highly shocked Shergotty and Zagami
meteorites (Duke, I968; McSween and Stolper,
I978; Stolper and McSween, I978; Smith andHervig, ~978) are characterized by about 7o~opyroxene (pigeonite En70Wo 11-En2owo2o andaugite En48Wo32-En33Wo32), c.2O~o maskelynite(An56Or3-An41Or3), c.2~ Ti-magnetite, c.o.5
ferrian ilmenite, and c.4 ~o mesostasis of fayalite,
silica, and K,Na, Ca-feldspars. Melting experiments
indicate approximately
one-third
cumuluspyroxene and two-thirds intercumulus liquid
formed at c. 141 o K. The shergottites are notable forthe intermediate composition of the plagioclase
glass, which contrasts with the oligoclase and albite
in ordinary and enstatite chondrites and the calcic
plagioclase in the eucrites, howardites, and C3
meteorites. Other volatile elements besides Na havea similar intermediate relationship between chondrites and eucrites. The shergottites are also distinguished from the eucrites by a higher oxidation
state, and the magnetite ilmenite assemblage equilibrated near the fayalite-silica-magnetite buffer.
Minor pyrrhotine (Smith and Hervig, ~978) andmerrillite occur in and near the mesostasis.
The eucrites also contain ferropigeonite andferroaugite but the plagioclase is very calcic(An85_~oo). The unbrecciated eucrites have igneous
textures, and the Ibitira specimen even has vesicles(Wilkening and Anders, ~975), which survived
later metamorphic events (Steele and Smith,I976). The Moore County (Hess and Henderson,
I949), Moama (Lovering, I975), and Serra de M/tge(Prinz et al., 1977c) eucrites have cumulate texture,and the Moore County eucrite has pyroxene assem-

73

blages indicative of slow cooling (Hostetler andDrake, I978 ). The brecciated eucrites have complex
textures, but clasts display igneous textures. They
contain much more U and Th than the unbrecciated eucrites (Morgan and Lovering, ~973), which
are very depleted in large-ion-lithophile elements.
The Ibitira unbrecciated eucrite contains no mesostasis and is extremely depleted in alkalis.
From melting experiments Stolper (i977) explained most eucrites by low-pressure fractionationof pigeonite and plagioclase from liquids similar incomposition to the Sioux Co. and Juvinas eucrites.
The compositions of the Stannern and Ibitira
eucrites cannot be produced by fractionation ofliquids with compositions like those of the othereucrites. Stolper suggested that a primitive
undifferentiated body yielded batches of magma by
partial melting of olivine Fa 35, pigeonite En 65Wo 5,
plagioclase An94, Cr-spinel, and metal. Initial melts
yielded liquid compositions like those of Stannern,
Ibitira, and Sioux Co., while later melts after
exhaustion of plagioclase yielded Mg-rich liquids
responsible for plutonic crystallization of Mg-richpyroxenes and olivine found in howardites, diogenites, and mesosiderites. The cumulate eucrites
may have accumulated from liquid~ resulting fromextensive fractional crystallization of advanced
partial melts.
Minor amounts of olivine, silica minerals, lowNi metal, troilite, chromite, apatite, and merrillite
occur in eucrites, but systematic petrographic andmineralogic studies are needed.
The pyroxenes show complex exsolution textures from which Takeda et al. (I976) and Miyamoto and Takeda 0977) deduced that eucrites
crystallized from o.5 to i2 km depth, This range isplausible for a large planetesimal, but the accuracyof the deduction may be affected by a multi-stage
metamorphic history, as proposed for the Ibitira
eucrite (Steele and Smith, 1976).
In general the eucrites resemble lunar basalts(McCarthy et al., 1973) but the details rule out
derivation from the Moon. The simplest assumption that the eucrites come from the same parent
body has allowed estimates of bulk composition
from several lines of geochemical reasoning.
Details vary, but most features are similar (e.g.
Anders, ~977; Consolmagno and Drake, I977;
Morgan et al., I978a). The bulk composition in
Table IV (Morgan et al., I978a) resembles that for
H-chondrites except for the lower content ofmoderately volatile elements Na, K, and Rb in
eucrites, and the lower Fe,Ni metal in eucrites,
though total Fe is similar. Perhaps the parent bodyof the shergottites is closer to H-chondrite thaneucrites because of the greater abundance ofalkalis.

74

J. V. SMITH

Diogenites. T h e nine hypersthene achondrites(Mason, I963b; one Antarctic meteorite) are composed essentially of orthopyroxene Fs23_2s withaccessory bytownite, olivine, troilite, chromite,tridymite, and metal. Six are monomict breccias,
Tatahouine is unbrecciated, and Garland andAioun E1 Atrouss are polymict. Their chemical
properties are quite similar, and cosmic-ray
exposure ages (Herzog and Cressy, i977) suggestthat five originated in a collision at - 14 x IO6
and perhaps three at c . - 24 x io 6 yr.
The Johnstown diogenite contains coarse clasts
in a fine matrix. Orthopyroxene EnTz_74Fs23_25
Wo2_3 occurs in fractured grains up to 2 cm across,some with triple junctions. Minor phases are
bytownite Ansz_9oOro.l, clinopyroxene En4w46
Fs9-10Wo44-45, olivine Fa28-29, silica mineral,
troilite, metal, and chromite. Floran et al. (I977b)
explained these data by impact deformation andrecrystallization of a primary cumulate. Gooley
and Moore 0976) found that metal in eight
diogenites occurs with chromite and troilite asarrays of inclusions predating fractures in orthopyroxene, and as larger grains in matrix. Most
metal has relatively low Ni (usually 0-5 wt ~o) andrelatively high Co (mostly o-7 wt ~o),but taenite (Ni15 53 wt ~o) occurs in three diogenites. A complex
origin is indicated: reduction of pyroxene with
migration of siderophile elements and silica tofractures and grain boundaries; annealing andrecrystallization of pyroxene with formation of
new grain boundaries; and possible addition of
metal from an external source. Miyamoto et al.
(I975) deduced slow cooling from the Mg, Feordering of Johnstown orthopyroxene, and notedthat the Ca content was too low to give hightemperature exsolution.
McCarthy et al. (I973) emphasized the association of eucrite and diogenite clasts in the polymict
brecciated howardites (next section), and Stolper's
model for eucrites involves partial melting ofolivine Fa35 , pigeonite En65Wos, and plagioclase
An94, Cr-spinel, and metal, all of which minerals
are represented in diogenites. However, detailed
studies are needed to test whether the plagioclase indiogenites (bytownite) is too sodic to yield eucritic
liquid by partial melting. Perhaps one planetesimal
supplied both the eucrites and diogenites, but there
were separate reservoirs on it. Radial zoning might
have been produced during accretion.
H o w a r d i t e s and mesosiderites. Both groups of
meteorites have complex unequilibrated mixturesof mineral and rock clasts, which have undergone
varying degrees of metamorphism both before andafter final agglomeration. The twenty mesosiderites
are approximately equal mixtures of silicates andmetal plus troilite, whereas the nineteen howardites

have only a little metal. Texturally and mineralogically the howardites match regolith breccias
from the Moon (Bunch, I975; Prinz et al., I977a)
but chemical compositions rule out derivationfrom the Moon. Although eucrite and hyperstheneachondrite (diogenite) clasts occur in both howardites and mesosiderites, the difference in other clast
populations suggests derivation from different
regoliths (Floran and Prinz, I978), perhaps on
different planetesimals formed after separate collisions with a single eucrite-diogenite planetesimal,or perhaps from more than one eucrite-diogenite
planetesimal.
Duke and Silver (I967) described the basic
mineralogy and petrography of howardites, andBunch (I975) gave a survey of the clasts, which are
dominated by pyroxenites and basaltic to gabbroicrock fragments resembling diogenites and eucrites,
but which also consist of types not hitherto recognized in meteorites.
The range of rock types is so similar to that inlunar breccias that details need not be given here;
let it suffice (Bunch, I975), that the howardites musthave originated in a body that differentiated thesame suite of rocks as the Moon. Clasts of CII
meteorites were observed by Bunch, Wilkening,and Dymek et al. The Kapoeta howardite (Dymek
et al., ~976) is rich in rock clasts set in a glassy
matrix comprised of comminuted mineral grains.
The petrographic features require plutonic-to-volcanic differentiation of basaltic material dominatedby pyroxene prior to thermal and impact metamorphism in a regolith. Dymek et al., concluded
that this howardite is not a simple mixture of
eucrite and diogenite, and noted that the Fe/Mn
ratio rules out derivation from the Moon. Crystallization ages demonstrated differentiation atc. - 4-55 to - 4.6 x Io 9 yr. The Malvern howardite(Desnoyers and Jerome, I977) has an unusuallylarge amount of glass (2o ~o). Hewins and Klein(I978) found that the Ni,Co pattern of the metal
has too little Co to match metals equilibrated
with silicates, and suggested that it results from'meteoritic' metal impacting the regolith.
Powell (I969, 1971) interpreted the mesosiderites
as unequilibrated polymict breccias because of thecomplex textures and the occurrence of Mg-richolivine with tridymite. The non-metallic portion isdominated by orthopyroxene(Fs20_40), pigeonite(Fs40_50W010-13), plagioclase (An80_98), and chromite. Rare metal fragments have Widmanst/itten
texture, but most kamacite nucleated at irregulargrain boundaries. The inferred cooling rate of o.I
K/IO 6 yr is even slower than for pallasites. Mason
and Jarosewich 0973) recorded a range ofpyroxenes from FS18_26Wo 2 (Veramin) to Fs35 52
W05 (Dyarrl Island) and noted an inverseM I N E R A L O G Y OF THE PLANETS
relation between bulk wt ~o and vol ~o of metal.
Nehru et al. (I978) recorded the following components in the Emery mesosiderite: mineral clastsof pyroxene, olivine, and metal; rock clasts ofdiogenite, eucrite, and cumulate eucrite; recrystallized metal-silicate clasts; recrystallized clasts of all
the above; and a norite clast. Floran et al. (I978)concluded that an impact-melt was the source of atleast the silicate portion of the Simondium, Hainholz, and Pinnaroo mesosiderites, and suggested
that the impact target was pyroxene-rich from the
preponderance of low-Ca pyroxene clasts. The
abundance of anorthite clasts was ascribed tocumulate eucrite and basaltic eucrite. Finally,
Floran and Prinz (1978) classified the mesosiderites, and concluded that silicate clasts strongly
dominate over rock and metal clasts, that silicate
clasts are dominated by low-Ca pyroxene (orthopyroxene and pigeonite), plagioclase and olivine,that diogenite is the dominant rock clast while
cumulate eucrite is abundant, and that there is arange of recrystallization textures.
Although the mesosiderites must have developedin a shocked regolith, they were released from theregolith by an event that did not shock themgreatly, judging from the preservation of inter-clast
textures.
The mesosiderites have a common origin (Mason
and Jarosewich, I973) because of sequential
composition range of pyroxene, near-constant
Ca, A1 ratio, and near-uniformity of trace elements
and Ni in the metal. Wasson et al. (I974) found thatthe metal in seventeen mesosiderites (Ni 7-9 wt ~o;
Ga ~3-16 Ge 47-58 Ir 2.4-4. 4 ppm) is in the samegeneral range as IIIAB and IIIE irons, pallasites
and H-chondrites, but does not fit exactly with anygroup. Weekeroo-type irons have higher Ni, Ga,
Ge, and Ir on an extension of the mesosiderite
trend, but the feldspar is K-rich rather than Ca-richas in the mesosiderites.
The complex polymict nature of the howardites
and mesosiderites makes a fitting close to thismineralogic and petrographic survey of the meteorites. There can be no doubt that eucrite anddiogenite fragments were reworked in one o r moreplanetary regoliths, and were mixed with othermaterial to give the howardites and mesosiderites,
but many uncertainties remain, including the possibility that the eucrite and diogenite fragments
differ in detail from the eucrite and diogenite
meteorites.
Oxygen isotopes. The ratios of the three oxygenisotopes constrain the genetic relations between the
Earth, Moon, and meteorites (Clayton andMayeda, 1978 ). The relations in fig. IX can be
explained only by invoking both mass fractionation between different phases reacting with the

75same reservoir of oxygen, and mechanical mixingof material from at least two reservoirs.
Simple mass fractionation produces phases
joined by a line of slope near o.5, such as the heavyline representing materials from the Earth andMoon, which apparently derived their oxygen fromthe same reservoir or the same mixture of reservoirs. A solid phase condensing from the solar
nebula below c.700 K is isotopically heavier than acoexisting gas phase (inset diagram), and the isotopic separation is greater at lower temperature(LT) than at higher temperature (HT). Mathematically the C2, Bencubbin, and Kakangari meteorites
could be related by mass fractionation, but not withany other specimens in fig. I I.
Rocks from the upper mantle of the Earth fall at
M, the position for the Moon. The oxygen isotopic
ratio of the Earth's lower mantle is unknown, andwill be assumed to be identical with that of theupper mantle. Equilibrium crystal-liquid differentiation of the whole Earth might give a lower
mantle that is isotopically heavier or lighter thanpoint M on the heavy line.
Since the meteorites do not lie on a single massfractionation line, mixing from more than one
reservoir is needed. The range for C3 meteorites
defines a line of slope I which is explainable byaddition of a minor component rich in 160, and thesimplest assumption is that the solar nebula wasinhomogeneous, probably because of addition ofvariable amounts of supernova debris rich in 160.
The meteorites can then be related by means ofmass-fractionation from gas-liquid-solid processes coupled with addition of ~60.
For this model the order of increasing additionof 160 is:(i) L and LL chondrites;(ii) H and CI chondrites; silicate inclusions inl i E irons;(iii) nahklites;(iv) E chondrites and achondrites; Moon andouter part of Earth;(v) eucrites, howardites, diogenites, most mesosiderites, two main groups of pallasites,
silicate inclusions in IAB irons, and Winona;(vi) ureilites;(vii) C2 and CK chondrites; Bencubbin andWeatherford stony-iron breccias;(viii) C3 and C4 chondrites;(ix) Eagle Station and Itzawisis pallasites.
Varying degrees of gas-solid, liquid-solid, or
solid-solid mass fractionation can be deduced. The
carbonaceous meteorites are isotopically heavier
than other meteorites because of condensation atlower temperature of solid from the residual gaseous nebula. Solid-liquid or solid-solid fractiona-

76

J. V. S M I T H

I'1'1

.

- -

L T

.

-

-

I0
8
6
0

-

r
L T~ ;-~
t/"of16 0

/

L/M

-J

4

u0
2
0

~ "

- -

-- /
/
-

-2

-4

M/.p

/

~_

/

'~ 2 ~ /

,/~..

-""
~
/4~"

Cl,C2,C3 Carbonaceous
Choedrites

HIL '

lAB

E
A,M,P

~ ,,~,,-~//
B e'K///-6~0/-/

-4

l,l,

-2

o

t

l,]i

4

6

s

MesosideritesandPollosites
NokhlaandLafayette

N,L
U
IAB,IIE~S

Ureilites
l~s

M

Moo~

K

Kokongori

S~80 rel. SM0W

-

it

14

16

-

_
--

--

BencubbinWeotherford,
EagleStation,Itzowisis

lll,lil,l10--~dieory Ch0ndrltes

EnstotiteChondrilesandAubriles
Euchres,
ilowordiies,Oiogenites,

i8

l-

20

Vig.H

FIG. I I. Oxygen isotopic ratios of meteorites referred to SMOW standard. The heavy line is the massfractionation line for earth materials (Clayton, I978).
tion at high temperature could explain the small
spread between the achondrites in group (v), butother evidence suggests that more than one
planetesimal is involved. The silicate inclusions in
IAB irons are too heavy isotopically to be related toany of the group (v) achondrites, and a different
planetesimal is needed. Clayton and Mayeda (1978)
suggested that the IAB silicates might be related to
the E-meteorites by addition of 160, as might Hchondrites to L and LL chondrites. The silicate
inclusions in liE irons are close enough to H
chondrites to be explainable by mass fractionation.
These isotope data reinforce the mineralogic andpetrographic evidence for a complex history of
meteorite formation, and conclusively demonstrate
the need for many parent bodies condensing from achemically inhomogeneous nebula. Although Emeteorites provide the closest match of oxygenisotopes to the Earth and Moon, so many mixing
models are possible (e.g. of ordinary chondrites+ Cz chondrite + basaltic achondrites) that there islittle isotopic control over the relation of meteorites
to the Earth, except to rule out more than small
amounts of meteorites with extreme compositions(e.g. all the carbonaceous meteorites).
Original location of the meteorites. That most
meteorites underwent prolonged metamorphism
and crystal liquid differentiation, together withlong exposure to cosmic rays, and that many passed
through a regolith, argues strongly for derivationfrom substantial planetesimals whose relics are
now represented by asteroids (Anders, I975). Only
the Apollo and Amor asteroids are currently reaching into Earth-crossing orbit, but these must contain material migrating from the main-belt
asteroids, mostly too small to be seen by telescopes.
Although incomplete matching of meteorite andasteroid spectra can be explained in many ways,
including the two possibilities that meteorites
represent only a tiny fraction of the diverse range ofasteroidal materials, and that asteroidal surfaces
undergo mineral sorting or selective degradation(e.g. Anders, I978), there is considerable concernthat the spectra of main-belt asteroids do notprovide a match for the most abundant meteorites--the ordinary chondrites. A host of differentiated silica-rich materials is required to matchthe iron meteorites, which themselves may represent only a small fraction of the M asteroids.
Some micrometeorites must come from the tailsof comets, but there is no concrete evidence thatany macrometeorites come from comets. It is
possible to envisage ejection by Jupiter of planetesireal fragments from the inner solar system andsubsequent transport by comets back to Earthintersecting orbit after a complex history in theM I N E R A L O G Y OF THE PLANETS
outer solar system. Comets may transport material
that originally condensed in the outer solar system.
It may be easier to invent ways of deriving carbonaceous meteorites from comets than the crystalliquid differentiated meteorites, but one must not77

overlook the strong evidence for polymict textures
and aqueous alteration. The next section willassume that all differentiated meteorites result
from condensation of the solar nebula (gas plus
solid particles) inside the orbit of Jupiter.

SUGGESTED SYNTHESIS
So many worlds, so much to do
So little done, such things to be.
Alfred, Lord Tennyson, In Memoriam
Whereas the preceding sections have tended torecite and criticize the available data and ideas onindividual planets, the present section attempts toprovide a comprehensive framework for further
studies.
Pre-planetary stage

To explain the oxygen isotopic ratios let us
assume that the nebula was a mixture of dust andgas, with some dust carrying supernova debris richin 160. The nebula did not undergo simple progressive condensation with equilibrium reaction ofsolids with residual gas. Some early condensates
were incorporated into substantial bodies, without
addition of more volatile components, as evidenced
by the irons rich in Ni and refractory siderophileelements. Many chemical types of planetesimals
were formed. There is abundant evidence of
crystal-liquid differentiation. Probably early
formed bodies were melted by short-lived radioactivity whereas late-formed ones stayed cool (timescale of IO6 yr for 26A1). Collisions of planetesimals
produced complex products including breccias
and chondrule-agglomerates. The planets wereassembled from many materials, which passed
through complex planetesimal stages. Chemical
models based on simple progressive condensation
are inaccurate in detail, but essentially correct
because of the statistical averaging o f the components entering the planets.
A crucial question is the redox state of thenebula, which is controlled by the C/O ratio when
H is highly abundant. Meteorites demonstrate awide redox state from the super-reduced enstatite
meteorites and Si-bearing irons, via a wide range ofreduced irons and stones with troilite and low-Fe
silicates, to the oxidized C I meteorites with magnetite. Let us assume that the oxidized meteorites
come from a cool part of the nebula in whichhydroxylated and hydrated silicates could form; tobe more precise, let us assign them mainly to theouter part of the main-belt asteroid region, andperhaps partly to the regi-9ns between the giant
planets. Then, the reduced and super-reduced
meteorites must be assigned to the inner solar

system, probably reaching to the inner realm of themain-belt asteroids. If so, how can Earth and Venus
have oxidized surfaces? Let us assume that the
inner planets began their accretion almost entirely
from reduced material in near-circular orbits (fig.
7), and ended their accretion with oxidized material
projected into the inner solar system by the lategrowing outer planets. The larger bodies (Earthand Venus) could have captured essentially all thehigh-speed oxidized material whereas the smaller
bodies (Mars, Mercury, and Moon) would havecaptured very little. Finally, the planets could havelost volatiles through the atmosphere and ionosphere, and even the Earth might have lost mostof its primordial H with concomitant oxidation ofC, S, and Fe. These assumptions of heterogeneous
accretion and planetary gas loss allow the Earthand Venus to have begun accretion from reduced,
and perhaps super-reduced material. Almost theentire population of planetesimals inside the orbitof Mars would have been captured by the inner
planets, but a few survivors might have reached themain belt of asteroids by a series of lucky encounters. This interpretation of the redox state
implies a radial gradient of C/O from >~ I in the
inner solar system (near Mercury and Venus?) to< I (near o.6? in the middle to outer solar system;
outside the Martian orbit?). But how was thisaccomplished? Did it involve grains of graphite in
the primordial nebula? Could radial migration ofgas and solids have been important (propulsion bysolar wind or radiation pressure?)? The present
synthesis will fail if answers do not appear to thesequestions.
There must have been some outward gradient ofpressure and temperature in the solar nebula, but it
seems desirable for the temperature to be low
enough for survival of isotopically unusual grains,probably derived from a super-nova. If this iscorrect the early lumps (fig. 3) could be a mixtureof surviving grains and new solar condensates.
Chemical equilibration would occur by melting
in early planetesimals heated by short-lived radioactivity rather than by solid-gas reactions in the
nebula.

78

J. V. S M I T H

EarthThe following f r a m e w o r k is suggested for a newmodel o f heterogeneous accretion (fig. 12). G r o w t h
b e g a n with accretion o f hot reduced material fromnear-by planetesimals, perhaps d o m i n a t e d bysuper-reduced enstatite-rich materials, as suggested
by naive interpretations o f oxygen isotopic ratios(fig. I 0. C o n t i n u o u s crystal-liquid differentiation
resulted in separation o f an Fe,Ni,S,C-rich core(including some K, U, a n d Si? ?) a n d a 'basaltic' rim(sensu lato) from a pyroxene-rich mantle. T h egrowing E a r t h would be s u r r o u n d e d by a thickcloud o f collision debris a n d gas.
Accretion c o n t i n u e d with expansion o f the feed. :.' :. ::'::.""ii:.'"),:-:)))idisintegrative ~":.'?."~ capture? ~-~(,db, Z]~
im OCt ~' P -."~ ~~176
coud
induced , o "Z]% afission?*'.-:,,."~*.
,o o impacting......
x ~planetesimaloasoltlc \ ~ [i?~:';~__~-valco,nic ~ ! ~ ]"nearly full grown
Moon near Roche limit
muchlost to Earth ~oil debris
captured~~(a)! ~ 0~
fully?

/ liquid\ ! Fe,Ni,S,etc.
Vcore (U,'K present ?)I0% grown

oxidation state increased chemistry strongly
by loss of H, especially a~.~;...~ h., ,..^

late ploneTesimols
more oxidized
M/Si >l,i/hardening,I ,J/montle j~conversionto

ao,,.,,, \,, ,, i Fe,S odies\ I I ~/ sinking throughM/ISi ~
reduced silicate\

ing zone as the growing E a r t h ' p u m p e d up' therelative velocity of u n c a p t u r e d planetesimals. Because o f the greater c o m p e t i t i o n from Venus t h a nMars, the c a p t u r e d planetesimals would tend withtime to c o m e increasingly f r o m the M a r s side with achemical tendency for higher redox state. Volatileloss from colliding planetesimals would tend toconvert F e - p o o r pyroxene a n d Fe-rich metal into
Fe-bearing olivine. P e r h a p s m o s t o f the last quarterof accreting material consisted o f differentiatedstony-iron planetesimals, represented in meteorites
by the A, M, P g r o u p in fig. I I, while a smalla m o u n t would consist o f material like ordinarychondrites, a n d a n even smaller a m o u n t would beof c a r b o n a c e o u s type, p e r h a p s closer in m e a n
c o m p o s i t i o n to C2 t h a n C I meteorites. The p r o p o r -.............
.r,h 0,o..
M/Si~Iliquid\/(b)

85% grown

FIG. 12. Cartoon of growth of Earth and Moon. (a) Early stage in which the baby Earth has g/own from hot bodies in the
middle of the Earth growth zone (fig. 7). These bodies are assumed to resemble enstatite meteorites, and havedifferentiated into a liquid core rich in Fe, Ni, S, etc. (perhaps with U, Th, and K?; Smith, I978) and an enstatite-richexterior through which negative diapirs of Fe,Ni,S-rich material are sinking. Strong volcanic activity is present at thesurface. A baby Moon may have grown by disintegration of an incoming planetesimal either with or without collision
with the Earth. A dust/gas cloud surrounds the Earth, and aerodynamic sorting tends to grade the impact debris fromfine particles at high altitude to heavy bodies near the Earth. (b) Late stage in which the Earth is now accreting
planetesimals that originally formed at considerable distance from the Earth, with a tendency to come from the Mars
rather than the Venus side. These planetesimals tend to be more oxidized and have a greater ratio of octahedral cations(M) to Si (i.e. they contain substantial olivine). The core is fully liquid, and the inner solid core will not develop for
another cAo 9 yr when sufficient cooling has occurred. The lower mantle does not equilibrate completely with theincoming material and is beginning to rigidity. Debris from Earth-colliding planetesimals is captured by the Earth androtational fission does not occur. Much of the debris from collision with the Moon ends up on Earth. Bodies perturbed
by Jupiter arrive at such a high velocity that the Moon retains very little of the collision debris. (c) Present Earthsegregated into a solid inner core especially rich in Ni and siderophile elements, a liquid outer core rich in S, P, C, andboth chalcophile and siderophile elements, a solid mantle zoned from pyroxene-rich chemistry in the lower mantle toolivine-rich chemistry in the upper mantle, and a complex crust containing most of the large-ion-lithophile elements.
The redox state of the mantle inherits the trend of the accreting planetesimals, and is modified by high-pressure effects in
Fe-bearing minerals. An extreme model has (Mg + Fe)/Si ranging from c. I to c.2 as Fe/Mg increases from c.o. I to 0.25
and then falls to c.o. I in the uppermost mantle. Convective overturning would modify this simple trend (which assumes
merely gravitational separation of Fe-rich liquid from heavier Mg-rich crystals), and the Fe/Mg variation may beconsiderably different. The uppermost mantle has 'pyrolite' bulk composition and consists of garnet and spinel
peridotites with minor eclogite, harzburgite, and dunite. Plate-tectonic processes would cause major local variations. In
the depth range 3oo-65 o km, several mineralogical transitions occur including the incorporation of pyroxene into
garnet solid solution and olivine into fl-phase. Below 65o km perovskite is the dominant structure type (perhaps in morethan one chemical variety), but other structure types including ilmenite should be present. In the lowermost mantle astructure (fluorite type?) denser than perovskite may occur.

M I N E R A L O G Y OF THE P L A N E T S
tions of the different types of accreting materialmust satisfy the oxygen isotopic ratios.
As the pressure of the lower mantle passed Ioo2oo kb (c.2ooo to 3ooo km radius), low-pressure
silicates typical of peridotitic compositions would
begin to transform into high-pressure silicates oroxides, unless the growing Earth became totallymolten from a catastrophic event. A growing
mantle in a relatively cool Earth would increasingly
hinder the sinking of Fe, Ni,S-differentiate fromaccreting material. Then the outer 5oo km of thepresent Earth would be able to retain some chemical features of the late-accreting material. Volcanic
activity and crust-mantle recycling would not
directly involve the lower mantle which would be
below the melting point because of the pressure
increase caused by subsequent accretion. Negative
diapirs of Fe,S-rich material would sink throughthe mantle without triggering off complete overturning. Loss of hydrogen would cause oxidation in
the outer part of the Earth. Some combination ofthese factors would allow the siderophile andchalcophile elements of the outer 5oo km to be'controlled by' oxidized iron in silicates coupledwith some depletion caused by sinking of sulphiderich liquid.
This model of cool differentiation of the laterstage of the Earth's growth does not preclude an
earlier history of hot differentiation. Indeed, thepossibility of collision-induced fission (fig. I2a) to
produce a proto-Moon must be considered during
growth of the first half of the Earth. Alternatively,
the Moon may have begun its growth near the
Roche stability limit from a massive debris cloudthat did not contain material torn off from the
Earth in a single catastrophic event. Whatever theorigin of the Moon it would accrete together with
the Earth after it became established in orbit. Many
complex processes would cause differences in thechemistry of material captured by the two bodies.
The Moon would be capturing only the lowvelocity fraction of the material, and would retain
very little of the volatile-rich material from theouter solar system.
The sequence of a hot early Earth and a cool late
Earth is plausible qualitatively from the energy
viewpoint only if three factors are valid: The earlyplanetesimals captured by the Earth must be hot
and the late planetesimals should be relatively cool.
Crystal-liquid differentiation should occur continuously during growth so that there is not asingle catastrophic overturning during the latestages of accretion. Finally, accretion must be slow
enough so that the gravitational energy of accretion can be dissipated sufficiently by radiation fromthe surface. Detailed calculations are needed toquantify these factors, and it must be emphasized

79

that the present model relies on chemical evidence
to claim that a mantle barrier existed during thefinal stage of accretion.
After completion of the major phase of accretion, volcanic activity at the surface would still be
intense, and would be triggered locally by the
decaying flux of the projectiles that formed basinson the Moon. Major impacts ceased by - 3-9 x ~o9
yr thereby allowing survival of some crustal relics
dating back to - 3 . 8 x I9 9 yr.
Many speculations are possible about the
Earth's atmosphere, but the key features for thepresent model are the major loss of H withconcomitant oxidation of C, S, and Fe in the outerpart of the Earth, and the establishment of a nonreducing and probably slightly oxidizing atmosphere by - 3.8 x 109.yr. Perhaps an intense solar
wind caused total dissipation of an early atmosphere, but pronounced loss of volatiles from hot
planetesimals must also be considered in models
for the distribution of the noble gases.
During the Archaean and Proterozoic eras thevolcanic processes would simmer down, and atsome stage plate tectonic events would develop.
However, these events did not cause multiple
cycling through the lower mantle. The only losses
from the lower mantle were gases and some othervolatile species, perhaps including alkali elements,
For the future there is the possibility that platetectonic processes might evolve into whole-mantle
convection in which a secondary differentiation
sweeps out material not involved in the primarydifferentiation. Such material might be dominatedby Na-Al-rich components, for which cosmochemically based models of the Earth's composition indicate substantial amounts in the mantle.
Finally, in the distant future, a cooling crustmight absorb most of the oceans and atmosphereby conversion of anhydrous minerals into hydroxylated, hydrated, and carbonated ones. Details are
difficult to predict because o f the effects of diffusion
and mechanical instability on the transportphenomena.
From the chemical viewpoint the present modelof heterogeneous accretion allows the possibility
that the lower mantle has a lower Fe content thanthe upper mantle because of the lower content of Ee
in silicates from reduced planetesimals than fromoxidized planetesimals and because of the concentration of Fe into rising liquid overlying Fepoor crystal cumulates; however, this simple
conclusion would be modified by gravitational
overturning before the lower mantle became too
rigid. The lower mantle would not have the same
composition as the upper mantle, and the garnetperidotite and pyrolite models would apply only tothe latter. A pyroxene-rich composition for the80

J. V. SMITH

lower mantle allows models for the bulk Earth inwhich the Mg/Si ratio is close to that for CI
meteorite. The Earth's core would be reduced
rather than oxidized, and the major light elements
would be S and perhaps C and P rather than O.

Other terrestrial planets
There is little to add for the Moon, except toemphasize that it may have started as a fission
product from the Earth (induced by disintegrative
glancing capture of a large planetesimal from theMars side of the Earth?), and proceeded to accrete
simultaneously with the Earth. The lower abundance of iron, siderophile elements, and volatileelements can be explained by a combination ofprocesses. Because the Earth and Moon havedifferent chemical compositions, it is necessary to
explain the near-identity of oxygen isotopic ratios
by fortuitous averaging of different mixtures ofaccreting material; perhaps the higher content of
C2-1ike material in the Earth is balanced by aconcomitant amount of material resembling
ordinary chondrites.
If this proposed synthesis is correct, Mercury
and Venus must have accreted reduced material
from the inner solar system, rather than theoxidized material assumed by earlier workers suchas Lewis.
Under reduced conditions Fe,Ni-rich alloy cancondense before Ca, A1,Ti-oxides and Mg-silicates(Table III), thus providing a simple explanation forthe high content of Fe in Mercury. However, thissimple idea is almost certainly complicated bymechanical fractionation of metal and silicate at
the planetesimal stage, especially because of thehigh collision velocities near the Sun. Furthermore, the Mercurian crust might have been partly
spalled off by late high-speed impacts. The bulk
composition in Table IV is based on oxidized
condensates, and should be replaced by one basedon reduced condensates. There is ample scope for
speculation on the detailed composition of the core,mantle, and crust of Mercury. Melting studies areneeded for bulk compositions in which the proportions of Ca, AI,Ti-oxides and Mg-silicates arechanged. The core should be low in S, and might
contain substantial amounts of C, P, N, andperhaps even Si; perhaps the Ni content is higher
than 5 ~o because of incomplete reaction of early
condensate (c.2o ~o Ni) with remaining gas. Low-Fe
pyroxene, spinel, and perovskite should be major
minerals in the mantle, and garnet may also be
important. Volcanic rocks should have unusual
low-Fe compositions undersaturated in silica.
Unfortunately the Fe-free minerals have featureless reflectance spectra, which will inhibit chemical

characterization by remote sensing. Furthermore,
the Mercurian surface may contain some Fe-richmaterial projected into the innermost solar system.
Venus may have had a similar history to the
Earth except for capture of all the orbiting debris(and probably a retrograde moon), and a higher
surface temperature. Perhaps Venus lost nearly allits hydrogen with concomitant oxidation of theatmosphere, crust, and perhaps outer mantle. However, the deep interior may be reduced. Although
the bulk composition of Venus may be generally
similar to that of the Earth, many speculations
about details are possible, including the amount of
S in the core, and the extent of metal-silicate
mechanical fractionation prior to accretion. Venus
should have received rather less Ca-like material
than the Earth, and measurements of greater
accuracy are needed to determine if it has a lower
K/U ratio. Because Venus is smaller than the Eartha greater fraction of the mantle occurs as lowpressure silicates rather than as high-pressure
phases. Sulphides should be absent in the crust andprobably in the upper mantle, and the sulphur
should occur at the surface as sulphates andperhaps as feldspathoids. Granitic and basaltic
rocks should be less abundant on Venus than on
Earth. The mantle should contain less Fe than theterrestrial one if the planetesimals in the Venusianzone were more reduced than in the terrestrial
zone.
Predictions for Mars are very uncertain because
it may lie near the transition zone from reduced tooxidized planetesimals, and because it is strongly
affected by gravitational effects from Jupiter. The
present main-belt asteroids are merely the remnants of a larger population, and at least someasteroids have moved either towards or away fromthe Sun since initial aggregation. I assume that theover-all distinction between an inner region dominated by Silicaceous or Stony-Iron asteroids andan outer region dominated by Carbonaceous(?)
ones is a relic of primordial accretion, while theunusual asteroids (e.g. E asteroids and Vesta) havechanged position. Perhaps the E asteroids originally formed inside the orbit of Mars and were
perturbed by Earth and Venus. If the above
assumption is correct, Mars accreted mainly from
S-asteroids, whose assumed stony-iron nature
should indicate early melting, destruction,
agglomeration, and metamorphism based on theevidence from stony-iron meteorites. If so, Mars
should have a fairly reduced interior since thestony-iron meteorites contain no ferric iron, andcontain sulphide rather than sulphate (the sulphideis mainly troilite). Mars should have accreted asmall amount of material from the outer part of themain belt of asteroids, and detailed mineralogical

MINERALOGY OF THE PLANETS
identification of the C-asteroids is badly needed.
For the present I assume that the C-asteroids tend
to be closer to C2 and C3 meteorites than to thehighly oxidized CI types, which are tentatively
assigned to the outer solar system. The oxidation
state of the Martian crust may be governed by a
combination of loss of hydrogen and of inheritanceof the oxidation state of C-asteroids. This assumes
that Mars did not become completely molten, or
did not melt sufficiently for the whole body toequilibrate. Perhaps the lavas forming the huge
volcanoes are reduced, except for the surface, which
has reacted with the oxidizing atmosphere. This
framework is opposed to the concept of a highly
oxidized Mars with a magnetite core. Rather, I
would prefer to explore the idea of a core rich in Fe,
Ni, and S, and a mantle with a lower average Fe
content than envisaged by other scientists. For apartly reduced Mars the moment of inertia would
be lower (c.o.37) than the present suggested value(0.377).
CONCLUSION
Looking forward to the twentieth century there
are plenty of opportunities for planetary exploration by the members of the Mineralogical Societyof Great Britain and Ireland. In spite of the great
advances of the past two decades there are many
uncertainties in our understanding even of the
Earth and Moon. Exploration techniques are now
developed sufficiently to allow determination of thechemistry, mineralogy, and petrology of the surfaces of all the terrestrial planets, but budget
restrictions will limit the rate of advance. Badly
needed to control theoretical speculation are seismic velocity and density profiles as well as accurate
values of the moments-of-inertia. Perhaps thegreatest mineralogical advances in the next fifty
years will come from coupled study of meteorites
and asteroids, while a new era will open up fromclose-range studies of the satellites of the giant
planets, for which long-range studies indicate
remarkable properties (e.g. alkali emission fromIo).
But no matter how many observations are mademost of the "detailed information about the earlysolar system has disappeared. Furthermore, there
seems little hope of obtaining definitive answers to
the depth profiles of minor and trace elements inplanets. We must he content with plausible answers

81

to many questions, though we need not be as
pessimistic as Ecclesiastes: 'That which is far off
and exceeding deep, who can find it out?' We canfeel confident that we are slowly working towards
an over-all understanding of the development of thesolar system, and that the great Architect has notplaced barriers to our scientific exploration of theprincipal features.
Returning to Paradise Lost, Book VIII, 66-84, let
us ponder how John Milton's words pertain to ourstudies:'To ask or search I blame thee not, for Heav'n
ls as the Book of God before thee set,
Wherein to read his wond'rous Works and learn
His Seasons, Hours, or Days, or Months, orYears:
This to attain, whether Heav'n move or Earth,
Imports not, if you reck'n right; the rest
From Man or Angel the great Architect
Did wisely to conceal, and not divulge
His secrets to be scann'd by them who ought
Rather admire; or if they list to try
Conjecture, he his Fabric of the Heav'ns
Hath left to thir disputes, perhaps to move
His laughter at thir quaint Opinions wide
Hereafter, when they come to model Heav'n
And calculate the Stars, how they will wield
The mighty frame, how build, unbuild, contrive
To save appearances, how gird the SphereWith Centric and Eccentric scribbl'd o'er,
Cycle and Epicycle, Orb in Orb.'

Acknowledgements. It is impossible to list all thosescientists who provided the basic knowledge used in thislecture, ranging from my mentors in crystallography andmineralogy at Cambridge University to my colleagues at
the Geophysical Laboratory, Pennsylvania State University, and the University of Chicago who taught me therudiments of geochemistryand meteoritics.A first draft ofthis lecture was criticizedconstructively by the following,
who must not be blamed for shortcomings in this revised
version: T. J. Ahrens, R. Brett, T. E. Bunch, R. N. Clayton,
M. J. Drake, D. M. Hunten, W. M. Kaula, B. Mason,
E. R. D. Scott, D. M. Shaw, S. C. Solomon, G. W.
Wetherill, and J. A. Wood. A particular debt is owed to
M. H. Hey for editorial and scientific advice. Irene
Baltuska typed the manuscripts with her usual skill, andBrenda Smith improved the style. During the past ten
years in which the present ideas simmered, financial support was provided by the National Aeronautics and Space
Administration and the National Science Foundation
for which current grants are NASA I4-Ooi -I7I and NSF
76 o36o49

82

J. V. S M I T H

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References
The following papers and books, which are relevant to the major topics
discussed in the Hallimond Lecture, were noticed between June and November
1978.
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Eloran (R.J.), Prinz (M.), Hlava (P.F.), Keil (K.), behru (C.E.) and
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OF THE

PLANETS89

Lebofsk,v (L.A.), 1977. Identification of water frost on Ca]listo.
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Loper (D.6.), 1978. The gravitationally powered dynamo. Geophys-.J. Roy.
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