lCARUS 34, 512-536 (1978)

The Mantle of Mars: Some Possible Geological Implications of Its High Density

T. R. M c G E T C H I N

Lunar and Planetary Institute, Houston, Texas 77058

AND

J. R. S M Y T H

Geosciences Division, Los Alamos Scientific Laboratory, University of California, Los Alamos, New Mexico 875~5

Received November 5, 1977; revised January 3, 1978

The density of the Mar t ian mantle is est imated to be about 3.55 g /cm 3 (Reasenberg, 1977). Model mineral assemblages for the Mar t ian mantle (at 30 kbar) were calculated using a modi- fied CIPW norm scheme by adding FeO to model terrestrial mantle compositions. The density of the resulting mineral assemblages vary with increasing FeO content. With pyrolite s tart ing compositions for the terrestrial mantle, the resulting model Mar t ian mant le with density of 3.55 g /cm 3 is not garnet-lherzolite like the Ear th ; rather it is an assemblage properly called oxide-garnet wehrlite : oxide (periclase-wiistite) 2% ; garnet 1 1 ~ ; olivine 73% ; clinopyroxene 1 2 ~ ; with no orthopyroxene. Part ia l melting of such an assemblage would yield iron-rich, ultrabasic lavas, with extremely low viscosities. Specifically, model partial melts, assuming production from the quaternary eutectic (inferred to be near: OpT g4~ cpx43 OXs) yields an ultrabasic (SiO.2, 41 to 44%) picritic alkMi-basaltic melt (norm composition ne 2.5, plag 32, or 2.4, di 20, ol 37, mt 4.4 and ilm, tr), with a computed viscosity of about 12 P at 1200°C. This model for the composition of the Mar t ian surface lavas (derived from geophysical data and petrologic arguments) is in remarkable agreement with a recently published model by Maderazzo and Huguenin (1977) (derived from reflection spectroscopy, experimental and theoretical nmdels h)r weathering in the Mar t ian environment) . The result also appears to be consistent with recent interpretat ions (Rasool arid Le Sergeant, 1977) of Viking atmospheric chemistry results, namely tha t the Mar t ian crust is potassium poor. There are a number of geological implications which follow, including (1) superfluid lavas may account for some flood and erosional features observed on Mars; (2) the X I l F inorganic chemistry experiment on Vikings 1 and 2 (Baird, 1976) indeed may be measuring compositions approaching primary lavas, contrary to current interpretat ions which favor a rather mature (weathered) soil; (3) ultrabasic (ferrokimberlitic) ash might be a major consti tuent of the Mar t ian soil, especially if cosmological models concerning the incorporation of a much volatile material within the early accreting Mars are correct - -a mat ter of current debate; (4) a number of mineral as- semblages riot previously considered are possible in the Mar t ian mantle depending principally on the activity of volatile substances, (S, O, C, H) ; i t is possible tha t some very unusual magmas are produced on partial melting; and (5) some fcrro-granite melts might be produced by liquid immiscibility.

1. INTRODUCTION significantly higher than those for the Current estimates of the density of the Ear th (3.3 to 3.4 g/era3). This difference

Mar t ian mantle (3.5 to 3.7 g / cm ~) are ahnost certainly implies a more iron-rich

0019-1035/78/0343-0512$02.00/0 Copyright ~ 1978 by Academic Press, Inc. All rights of reproduction in any form reserved.

512

IMPLICATIONS OF MARS MANTLE DENSITY 513

mantle composition, probably present as FeO (possibly as FeS and/or Fe20~ or Fe304). A qualitatively different mineralogy may be implied. In this paper we explore high pressure (~30 kbar) mineralogies which we believe to be petrologically plausible and consistent with the Mars mantle density. We find that, indeed, the model Martian mantle has a different phase assemblage than the Earth and that partial melting of this mantle would be likely to produce ultrabasie (ferrobasaltic) melts of very low viscosity.

We believe the implications, which are consistent with geology and geophysics (from Mariner 4, 8, and 9 and also with preliminary results from Viking 1 and 2), include the following : (1) very low viscosity lavas, probably ultrabasic in composition rather than basic, may be common on Mars, (2) this low viscosity is consistent with large eruption rates producing copious floods which may have been a dominant process in shaping the Martian surface, (3) volcanic ash (ferro-kimberlite) could be an abundant constituent of the soil-- especially if volatiles are abundant in the Martian interior as predicted by some currently popular cosmological models, (4) stream and channel deposits commonly attributed to water erosion may be lava- erosion features, (5) terrestrial analogs may not exist for either surface processes nor

TABLE I P R O P E R T I E S OF THE E A R T H AND M A R S

M a r s E a r t h

M e a n radius (km) 3390 .0 a 6378

Mass (gin) 6 . 4 1 8 M i02aa 5 . 9 7 6 X 102~

Bu lk densi ty (g/cma) 3 . 9 3 3 a 5. 517

Calcu la ted uncompressed densi ty (g/cm3) c 3 . 7 4 . 0 3

Dimensionless polar moment of iner t ia ( C / M R 2) 0. 365 b 0. 3335~

Mass f ract ion in core~ 0 . 1 9 0. 315

a Saunders (1976). b Reasenberg (1977).

K a u l a (1968, Tab le 9 . 5 ) .

the surface itself. Specifically, (a) Martian lava fluxes could greatly exceed rates ob- served in modern terrestrial eruptions and (b) the Martian surface could be inter- bedded, thin ultrabasic flows, kimberlitic ash, ice and wind-blown material. The response of such a surface to wind erosional processes, overriding by younger flows, intrusion by melts from below, slope- failure, and global extension are interesting grounds for speculation, observation, and analysis.

2. CONSTRAINTS ON THE COMPOSITION OF THE MARTIAN MANTLE

The Martian mantle (MM) density is constrained by the bulk density of the planet, its moment of inertia, and model assumptions of the size and density of the core. Although the mean uncompressed density of the planet is less than that of the Earth, probably implying a lower bulk iron content, its small core (relative to the size of the whole planet) suggests that the Martian mantle is more dense than the Earth's (see Table I). A recent detailed review by Johnston and Toks6z (1976) give solutions for the mantle density of Mars as functions of the parameters mentioned above. In Fig. 1, the Johnston-Toks6z (JT) solutions are shown as functions of the parameter C /MR s , the polar moment of inertia, derived from the motion of artificial satellites, such as the Mariner and Viking spacecraft. Reasenberg (1977) recently re- ported a value of 0.3654 ±0.001 for C / M R 2 from analysis of Mariner 9 tracking data, in which he includes a model assuming rigid support of the excess mass of the Tharsis ridge. The resulting value of C /MR s is substantially lower than previous estimates of 0.377 (Lorell et al., 1972; Johnston el al., 1974). This implies a some- what lower MM density than previously accepted, closer to 3.55 than 3.7 g/em 3, but still substantially higher than the 3.34 commonly accepted for the Earth. Okal and Anderson (1978) recently published a

514 McGETCHIN AND SMYTH

4,0

5.B F-

~_ 5.6

I i I

3.4

5.2 0.55

BAND of SOLUTIONS for CONDUCTION THERMAL MODELS

0 . 5 6 5 4 JT R J ' ~ ' ~ R EASEN BE RG , 7 ~ /

1977 ~ /

/ 0T, ~6 I I I

0 . 3 6 0 . 5 7 0 . 3 8

C/MR 2

FIG. 1. Dimensionless moment of inertia and the mantle density of Mars, for various assumed core compositions and internal heat-transfer mechanisms, after Johnston and Toks6z (1976).

Mars model using the Reasenberg mo- ment of inertia which has an upper mantle density of 3.40, believed to be a lower bound (Anderson, 1978, pers. comm.) with a higher value likely.

Factors proposed to account for the observed density differences among the terrestrial planets include (1) differing iron abundance, or more simply Si/Fe ratio (Urey, 1952), (2) differing oxidation- reduction states (differing Fe/Fe oxides and Si/SiO2) but generally similar Si/Fe ratios (Ringwood, 1966), or (3) differing sulfur and volatile contents, together with different FeO abundance (Lewis, 1972). Grossman and Larimer (1974) point out that differences in oxidation state alone and volatile content cannot account for the high density of Mercury and the low density of the Moon. Furthermore, the wide variability of Fe/Si ratio in chondritic meteorites argues strongly against an ad hoc assumption for Fe/Si uniformity. Urey (1952), on the basis of mean density of the planets, estimated that the Fe/Si ratio of Mercury is three times that of the Earth; five times that of Mars.

Models which assume condensation of material from the primitive solar nebula under thermodynamic equilibrium condi-

tions (e.g., Cameron, 1963; Cameron and Pine, 1973) in which P and T are deter- mined by heliocentric distance and time, produce bulk compositions in good agree- ment with observed planetary densities (Lewis, 1972, 1974; Grossman and Latimer, 1974; Clark et al., 1972; Anderson, 1972). In general these models predict that Mercury is iron-rich and anhydrous because it formed at temperatures so high that silicates were not totally condensed. While Venus and the Earth are similar, Venus contains no sulfur and no water--a pre- diction which is at variance with the recently reported H:S04 content of the Venus atmosphere by the Soviet Venera 4 experiment. These models predict a bulk Mars which is iron deficient and volatile- rich relative to the bulk Earth ; and a Mars mantle, which is more iron-rich than the Earth's, and probably more oxidized and volatile-rich. There is, however, substantial disagreement about the volatile contents of Mars. Recent publications (Anders and Owen, 1977a,b,c) argue for a Mars which is severely depleted relative to the Earth. Also, the Moon remains an obvious anomaly in the equilibrium condensation models (Anderson and Ringwood, 1977) and requires special treatment in order to be reconciled with density-heliocentric distance systematics.

The present oxidation state of the iron in the Martian mantle is a crucial and essentially unknowable property. The oxy- gen fugacity in the Earth's interior is not known with confidence nor whether the Earth's mantle is in chemical equilibrium with its metal-rich core or not (Ringwood, 1966; Brett, 1976). I t may be constant and low fo2 as implied by Brett (1976), or it could vary greatly from place to place (Ringwood, 1966). Brett (1976) has re- cently reviewed and summarized this argu- ment, in the context of the origin of the Earth's core. Sato (1968) has suggested that the fo~-T relations in terrestrial basalts are consistent with buffering by

IMPLICATIONS OF MARS MANTLE DENSITY

TABLE II

COMPOSITION OF THE MANTLES OF THE EARTH AND MARS

515

Earth s Undepleted Depleted

Mars b This paper

1 2 3 4 5 6 7

Mineralogy

Chemistry

olivine 56 67 45 opx 16 12 40 cpx 14 11 10 garnet 13 10 5 oxide other

SiO2 45.16 44.18 48.92 TiO2 0.71 0.09 0.07 A1203 3.54 2.81 1.97 Cr:O3 0.43 0.30 0.34 Fe:03 0.46 1.16 0.27 FeO 8.04 7.34 6.03 MnO 0.14 0.14 0.12 NiO 0.20 - - 0.25 BaO 0.01 - - 0.003 MgO 37.47 40.95 39.39 CaO 3.08 2.49 2.40 Na20 0.57 0.22 0.19 K20 0.13 0.04 <0.01 P205 0.06 - - 0.04

73

12 11 2 1

41.05

3.22

0.42 16.40

34.64 39.03 42.32

2.72 3.06 3.32

0.35 0.40 0.43 29.47 20.53 13.82

34.07 28.74 32.38 35.11 2.80 2.36 2.66 2.89

0.12 0.10 0.11 0.12

a (1) Pyrolite I I I (Green and Ringwood, 1963). (2) Mean of 5 high-calcium olivine nodules (Harris et al., 1967). (3) Garnet lherzolite nodule, Wesselton Mine, So. Africa (O'Hara, 1968).

b (4) Calculated Mars mantle composition, this paper. (5) Mars mantle (Johnston et al., 1974), C/MR s = 0. 377, p = 3.7. (6) Mars mantle (Johnston and Tolks5z, 1976), C/MR s = 0. 367; Core FessS16. (7) Mars mantle (Johnston and TolksSz, 1976), C/MR s = 0.365; Core FeS.

e l e m e n t a l c a r b o n (g raph i t e ) in t he man t l e . T h e s t a b i l i t y r e l a t ions in s u l f u r - w a t e r - oxygen b e a r i n g p e r i d o t i t e sy s t ems a t h igh pressure are c ruc ia l for u n d e r s t a n d i n g redox r e l a t ions wi th in the p l a n e t s b u t t h e y are no t well known. T h e r ecen t w o r k of P o p p et al. (1977) on a m p h i b o l e s t a b i l i t y c l ea r ly unde r sco re s t h e i r s ignif icance a n d the need for more e x p e r i m e n t a l w o r k in these s y s t e m s w i th M a r s in mind .

T h e cosmolog ica l m o d e l s g e n e r a l l y agree t h a t t h e m o r e r e f r a c t o r y ca t ions (A1, Mg , Ca, p o s s i b l y Si) were n o t s e v e r e l y f rac- t i o n a t e d d u r i n g acc re t i on processes in t h e v i c i n i t y of t h e E a r t h a n d M a r s .

R e g i o n a l geo logy and t h e geophys i c s of M a r s p r o v i d e some useful l im i t s on t h e

d e p t h of m a g m a genesis . T h e he igh t s of t h e l a rge shie ld vo lcanos a long t h e T h a r s i s r idge are e s t i m a t e d b y Blas ius and C u t t s (1976) to be over 20 k m w i t h t h e h e i g h t of M o n s O l y m p u s be ing a b o u t 24 km. F o r a d e n s i t y c o n t r a s t of a p p r o x i m a t e l y 1 0 % b e t w e e n t h e f luid m a g m a and a v e r a g e d e n s i t y of t h e rock c o l u m n t r a v e r s e d , h y d r o s t a t i c e q u i l i b r i u m d i c t a t e s a r e se rvo i r d e p t h of 150 to 250 k m (Car r , 1973, 1974), e q u i v a l e n t to 20 to 32 k b a r of l i t h o s t a t i c p ressu re w i t h i n M a r s . T h e s e d e p t h s a re c ons i s t e n t w i th e s t i m a t e s of t h e t h i cknes s of t h e M a r t i a n l i t h o s p h e r e r e q u i r e d to s u s t a i n t h e sh ie ld vo l canoes as s t a t i c l oads and also t h e d e p t h of in fe r red i so s t a t i c

c o m p e n s a t i o n r e q u i r e d to a c c o u n t for

516 MeGETCHIN AND SMYTH

T A B L E I I I

NORM MINERALS AND STP DENSITIES USED IN CALCULATIONS

Mineral Formula STP density

Pyroxenes

K-Jadeite KA1Si20 ~ 3. 330 Jadeite NaA1Si206 3.315 Acmite NaFeSi20~ 3. 750 Diopside CaMgSi20~ 3. 277 IIedenbergite CaFeSi~O6 3. 550 Enstatite Mg2Si206 3. 198 Ferrosilite Fe~Si~O 6 3. 966

Olivines

Forsterite Mg2SiO4 3. 214 Fayalite Fe2SiO4 4. 393

Garnets

Andradite Ca.~Fe2Si301~ 3. 860 Pyrope MgaA12SiaOr2 3. 559 Almandine Fe3A12Si3OI~ 4. 318

Oxides

Periclase MgO 3. 584 Wiistite FeO 5. 745 Ilmenite FeTiO,~ 4. 786

Bougier gravi ty low over the Tharsis ridge (Phillips et al., 1973).

The above arguments suggest tha t (1) the Mar t ian mantle has a density of approximately 3.55 g /cm 3, (2) this density implies Mars has an iron-rich mantle, relative to the Earth 's , (3) the depth of origin of the shield lavas is 150 to 250 km (approximately 20 to 30 kbar), and (4) volatiles are likely to be present during partial mel t ing--a l though this is current ly a topic of debate.

Rasool and Le Sergeant (1977) have pointed out tha t while the *~A/36A ratio observed by Viking is 10 times the value of the Earth, tha t the absolute amount of 4°A is very low (1 X 10 -4 tha t of ordinary chondrites, relative to the Ear th ' s value

of 4.1 X 10 -2, both normalized to carbon). They suggest tha t these differences m ay be accounted for by a low abundance of 4°K in the outer layers (lithosphere) of Mars and an absence of intense early out- gassing. They estimate values of K20 in surface materials near 0.25% K20, sub- stantially lower than the 1.6% average values for the Ear th ' s crust (but consistent with or even higher than fresh terrestrial midoeean ridge basalts.

3. COMPOSITION AND MINERAL ASSEM- BLAGES OF THE MARTIAN MANTLE

We wish to explore plausible mineral assemblages for Mars, at 30 kbar with chemical compositions consistent with a density of 3.55 g /cm 3.

The est imated tempera ture within Mars at 30- to 35-kbar pressure, tha t is, about 250-km depth, is 1100 to 1200°C (Johnston et al., 1974, Fig. 11), well above the stability of amphibole, perhaps the most likely hydrous mineral. Although phlogopite (Kushiro et al., 1967) and possibly humite might exist to this depth in minor amounts, we can reasonably restrict ourselves to anhydrous assemblages; in this view, any volatiles present would form a discrete phase.

The effects which can act to increase the density of any peridotite assemblage in- elude: (1) compressibility resulting from increased pressure- -a small effect, not applicable here; (2) decrease in tempera- t u re - - a s above; (3) iron enr ichment - -as FeO, FeS, or metal; (4) garnet enrichment - - a n interesting possibility especially if Fe 3+ is abundant in Mars; (5) very low fo~--result ing in metallic Fe; (6) high fo~--result ing in creation of magnet i te + opx at the expense of olivine, (or perhaps Fe 3+ garnet at high pressure) ; and (7) high fs~--resulting in an assemblage of sulfide, oxide, and opx at the expense of olivine, although the extreme upper bound on

Step

IMPLICATIONS OF MARS MANTLE DENSITY

Mineral Cation compositions

Resulting mineral phase

517

(1) Ti (2) K (3) Na

(4) Fe 3+ (5) Fe~+/Mg

(6) Ca

(7) A1

(8)

i f> 1 /z

Mg + Fe] ~tom ¼

i f < l

) FeTiO3 }ihnenite " ilm ) KA1Si206 ) ) epx~ ) NaA1Si206 ~"jadeitic" cpx

NaFe3+Si~O 6 Ca3Fe23+Si30,2 } andradite ) g~ ratio determined, partitioned to

ol, cpx, opx, g and oxide (K=I) ) CaMgSi2()6 "~diopside ) cpx~

CaFe~+Si'.,O 6 Jhedenbergite +(2) ) Mg~Al2Si~Ol~. ~pyrope , g~

Fe~2+A12Si30~: |almandine +(3) MgSiO3"~enstatite FeSiO3 Jferrosilite ) opx~

opx oliv ~

~Mg~SiO4) forsterite |Fe2SiO4 Jfayolite ' > oliv~

olivJ

°xide/MgO-~perielase [FeO Jwfistite " ) oxides,

Fro. 2. CMculation algorithm for high pressure norms showing the distribution of cations to mineral phases.

sulfur content would be about 7% assuming meteoritic (CC1) values.

Of the above, simple addition of FeO to a mantle composition which is otherwise Earthl ike has the vir tue of simplicity and is also consistent with (rather disparate) cosmological theory. In effect it says (1) the upper mantles of the terrestrial planets ( that is, Mars and the Ear th) have similar ratios of the major cations, except iron, a plausible assumption and (2) tha t the cation to oxygen ratios are similar. Also, as we shall see later, the activities of various species (S, O, H, C) will be very important .

We need to determine a model composi- t ion of the Ear th ' s mantle, in order to proceed. Ultramafie nodules in terrestrial basalts and kimberlites provide direct evidence of the composition and mineral assemblages within the Ea r th to several hundred kilometers depth. Most students of basalt genesis and of deep-seated nodules

agree tha t the Ear th ' s upper mant le con- sists predominant ly of magnesium-rich garnet- lherzoli te ( that is, a peridoti te con- sisting of garnet -t- opx + epx -t- olivine) with minor lenses and pods of eelogite (garnet + j adeitic epx). (The existing l i terature on this topic is too extensive to review here - -Yoder (1976) contains a good review of this argument.) Various estimates of the compositions of the Ear th ' s mant le are shown in Table II . A commonly accepted model is the hypothet ical mix of dunite and basalt, called "pyrolite" (Green and Ringwood, 1963, 1967a,b) which would (1) produce basaltic magma under proper partial melting conditions and (2) consti tute garnet-lherzolite under high pressure conditions. The principal problem with pyroli te is tha t at high pressure it would be very garnet rich, hence dense- - somewhat denser than most models of the Ear th ' s interior derived from inversion of geophysical data (e.g., Press, 1969, 1970)

518 M c G E T C H I N AND S M Y T H

3.7

to a

3..5

3.4

3.3

, - - I - - I - - I I I I I

3,55 i .

I I I I I I I I 0 4 s 12 16 20 24 28

w t FeO ADDED to PYROLITE TiT

FIG. 3. Calculated (STP) mantle density as a function of compositions consisting of pyrolite I I I plus varying amounts of FeO.

which lead to models of depleted or garnet- poor pyrolite. Recent seismological data (Sipkin and Jordon, 1976) and petrologic evidence (Boyd, 1976) suggest garnet-rich mantles, quite like pyrolite and unlike more depleted assemblages typical of garnet lherzolite nodules in kimberlites.

For our purpose, we assumed a terrestrial pyrolite mantle (pyrolite III) and pro- ceeded to calculate mineral assemblages for the Martian mantle as follows. Using a mantle-norm scheme (similar to CIPW norms) mineralogies of various hypothetical

chemical compositions were calculated by (1) adding FeO to pyrolite I I I in varying amounts, (2) calculating the relative abun- dance of the minerals resulting from each such composition, and (3) determining the (STP) density of the resulting assemblages. This is essentially the same procedure used previously by Johnston et al. (1974), how- ever we have amended their norm-algorithm considerably to be more consistent with what we believe to be plausible mineral- ogies. The principal difference is that we cast CaO into diopside, rather than garnet, because diopside is a major mineral con- stituent of deep-seated nodules.

Mineral densities used to calculate rock values are STP values, hence calculated rock densities will vary somewhat from their values in si tu (at. 30 kbar and approxi- mately 1100°C, the conditions at 250-kin depth within Mars). However, because of the compensatory effects of thermal ex- pansion and compressibility, this difference will be small (Sipkin and Jordon, 1976).

The list of mantle-norm minerals used in our calculations and their 30-kbar densities are shown in Table III.

The sequence of assignment of major cations to various mineral phases, is the basis for predicting the mineral assemblages of our hypothetical Martian mantle rocks.

I 0 0

90-

z

8 0 - -

~ 7 G - -

6O

o z

5O

I z / ' I ~ I 1.0 ILMENITE

GARNET ~ - o . 9

". "".cPx -0.8

o~* "'-_.._ - -i°'r= - ~

• I I • ~ /05 3 . 4 3 ,5 3 .6 3 . 7

C A L C U L A T E D ' D E N S I T Y

FIG. 4. Calculated normative abundance of minerals as a funct.ion of calculated (STP) density, for models consisting of pyrolite I I I plus FeO iu varying amounts. Also shown is the resulting M g / ( M g + Fe) a tom ratio in all marie minerals, for no partitioning, tha t is K~ ~ 1. Darts, marked E and M, indicate the densi¢ies of the mantles of Ear th and Mars.

IMPLICATIONS OF MARS MANTLE DENSITY 519

: ~: MOST NODULES

~ KIMBERLITE

~ / % I PLUS ONE ALUMINOUS ~ P H A S E

NORM PERIDOTITE CL'~ SSIFICATION

FIG. 5. Diagram showing the classification of peridotitic rocks, and the calculated compositions of the mantles of Earth (circle with cross) and Mars (circle with arrow). Rocks believed to be derived from the terrestrial mantle are garnet and spinel lherzolite. Calculated Mars mantle is not lherzolite; it is oxide-garnet wehrlite.

The general features of the algorithm are reproduced in Fig. 2. There is much implied in this s cheme- -many deviations from it are possible; some of the obvious possi- bilities among these are discussed later. The basic norm scheme makes the following operations and cation assignments (1) Ti02 to ilmenite, (2) the A1/Fe 3+ ratio deter-

mined, (3, 4) K and Na to jadeitic and acmitic clinopyroxene, (5) Fe a+ to andradite garnet, (6) Fe~+/Mg ratio determined and equally parti t ioned to several minerals all with K D - 1, (7) Ca to diopsidic clino- pyroxene, (8) A1, remaining after step 2, to pyrope-a lmandine garnet, and (9) re- maining Si, Mg, and Fe z+ distributed to either orthopyroxene + olivine or periclase

olivine, depending on the S i /Mg + Fe a tom ratio. An expanded algorithm treats CaO and A12Q-rich compositions but for compositions near peridotite, the above works in a straightforward way.

Densities of various calculated model mantle compositions as a function of FeO (added to pyrolite I I I ) are shown in Fig. 3. With about 12.8% added FeO, the density matches the Mars value implied by the currently accepted C / M R s of 0.365, namely 3.55 g / c m 3 (for an FeSs~Fel~ core, see Fig. 1). This implies a total FeO content of 16.4% for the Mart ian mantle (see Table I I) .

Model mineral assemblages are shown as a function of density in Fig. 4. The dart marked E is the calculated mineralogy for

TABLE IV

MINERAL ASSEMBLAGES OF CALCULATED MANTLES FOR EARTH AND MARS

Earth Mars Pyrolite III Pyrolite III -{- 12.8 FeO p = 3. 383 p = 3. 551

fPyr 72.3 Garnet 12.72 ~And -F Sp 17.4

I.Alm 10.3

I K-Jd 4.4 Jd 23.2

Cpx 13.70 Acre 4.5

L Hd 7.8 Di 6O. 1

En 86.7 Opx 15.94 Fs 13.3

Fo 85.6 Olivine 56.24 Fa 14.4

Ihnenite 1.35

Garnet 11.39

Cpx 12.15

Oxide 2.14

Olivine 73.15

Ihnenite 1.18

f Pyr 58.8 And -[- Sp 16.8 Alm 24.4

K-Jd 4.3 Jd 22.9 Ac 4.5 Hd 18.8 Di 49.5

Pr 62.6 Wu 37.4

Fo 67.3 Fa 32.7

5 2 0 M c G E T C H I N A N D S M Y T H

.<

Z

-5. ©

c ¸

• ~ ¢~

¢9

I

¢9

©

¢-,

a9

©

I

?

I

IMPLICATIONS OF MARS MANTLE DENSITY 521

~ ~ •

~:~ ~ " ~

~ . ~ ,~

"~ .~

i ~'~ ~

~:~'~

~ ~

~ + b • ~

-~ ~..~

o~ ~.~

~ + ~

oo

I

~ ?

I

"l ~

©

I

0

-8 2~

,,-.

i

~.~ ,~ . ~l,

~p

l

522 McGETCHIN AND SMYTH

(pyrolite III) the Earth; the dart marked M is the Mars value. Two important differences are evident: (1) orthopyroxene in the Earth's mantle (0 "" 3.35) is replaced by oxide in Mars (o ~ 3.55), and (2) the mineral phases are much more iron-rich in Mars.

So, whereas the terrestrial mantle com- position (pyrolite III) calculates (as it should) to garnet-lherzolite (garnet- olivine-opx-cpx), the Martian mantle is an oxide-bearing (periclase-wfistite) (ferro) exist within this bulk chemistry--some of the more obvious variants are listed in Table V. The stability relations among the postulated iron-rich phases are functions of P, T, and the fugacities of various volatile components--it is likely that the effects of changes in fo2 and fs~ will be especially interesting. Several very unusual rocks could result. For example: (1) if epx is unstable in the presence of perielase- wiistite, then calcium-rich olivines could be produced, perhaps involving an inter- mediate mineral phase such as merwinite, CaaMgSi04 (Roy, 1956), resulting in either two-olivine garnet wehrlite (presuming perielase is used up before epx), or (2) if garnet is unstable with oxide, it may react to form olivine plus spinel, producing an unusual spinel-rich wehrlite. These reac- tions are outside usual lherzolite assem- blages, although Padovani (pers. comm., 1978) has recently reported kornerupine- bearing nodules from a kimberlite pipe (Moses Rock, Utah) which may be such an assemblage. Many other interesting possibilities exist and suggest some po- tentially valuable experimental studies in anhydrous Fe rich systems.

It is known that the stability field of the olivines are strongly influenced by fo2 and, in fact, this field is fairly restricted (see Williams, 1971; Speidel and Nafziger, 1968; Nitsan, 1974; and the earlier work of Muan and Osborn, 1956). At a tempera- ture of 1000°C, olivine is stable only at fo= between -- 12 and -- 16; at higher values,

olivine decomposes according to the reaction

6Fe2SiO4 -t- 202 ~ 4Fe304 -t- 3SIO2 (fa) (mt) (qtz)

or probably

6Fe~SiO4 + 02 ~ 2Fe304 -t- 6FeSiO3 (fa) (mt) (fs)

at pressures above about 2 kbar. At low fo~ values, iron metal is stable.

In a sulfur-rich Mars, the following reaction would be expected to proceed to the right,

4Fe2SiQ + $2 ~- 2FeS -t- 2Fe304 ~- 4SIO2, (fa) (po) (rot) (qtz)

or for a peridotite assemblage a more likely reaction is,

4Fe2SiO4 + 4Mg2SiO4 -/- $2 ~- (fa) (fo)

2FeS q- 2Fe304 -1- 4Mg2Si206. (po) (mt) (en)

The position of the equilibrium would be determined by log fs~ which for terrestrial lavas has been determined to be in the range 0 to - 4 , (Toulmin and Barton, 1964; Carmiehael, 1964). The resulting assem- blage would be rich in magnetite sulfide, and pyroxene but olivine-poor. One would expect partial melts from such a mantle to be relatively sulfur-rich with unusual com- positions, although sulfur solubility in silicate melt of any plausible composition is unlikely to exceed about 4000 ppm (Brett, personal communication). In a Martian mantle with high fo2 and high fs2, scapolite (Ca4(A128i08)3(SO4)) may be a stable phase, possibly with hematite.

Prinn (1975) recently made the sugges- tion for Venus that hydrogen loss from magma may drive the following equilibrium toward the right, producing oxide at the expense to sulfide,

2FeS ~- l lH20 ~ Fe20~ -~- 2H2SO4 ~- 9H2.

IMPLICATIONS OF MARS MANTLE DENSITY 523

H2SO4 has not been observed in the Martian atmosphere, but this does not preclude the possibility of its consumption during chemical weathering, a process which could be very important at depths of perhaps 100 m to 10 km within Mars.

The addition of water to the Martian mantle would decrease the density some- what but not significantly, if the resulting phase is amphibole. Equilibrium condensa- tion models (e.g. Lewis, 1972) imply high water contents and water has an important effect on strength, physical properties, melting point, and chemistry of melts produced so the question is a very im- portant one. Iron-rich amphiboles have reasonably high densities for example cure- mingtonite (Mg, Fe~+)TSisO22(OH)2 with (Fe~+/Mg + Fe 2+) of 0.7 has an STP density of about 3.2 ; actinolite, the calcium bearing equivalent at a similar Mg/Fe ratio has a density of about 3.15 (see Deer et al., 1966). A range of amphibole com- positions might occur, by the following reactions

a n + d i + e n + H20 Ca2A12Mg3SisO22 (OH) ~,

(hornblende)

3en + 2di + qtz + H~O --~ Ca2MgsSisO22 (OH) 2,

(tremolite)

where "Mg" is Mg + Fe 2+. Because parga- site (and K-rich richterite) is observed in

ultramafic nodules (e.g., Francis, 1976), amphibole is a particularly likely candidate, and a likely site for sodium as well,

d i + 3 e n W a b + a n + H 2 0 - - ~

or

3di + 2fo + 2ne + an + sp + 2H20 --*

NaCa2Mg4AI~Si602~ (OH) 2 + 4SiO~ (pargasite)

Clearly a range of amphibole compositions can be postulated which could store some of the water implied by the cosmological models in the Martian mantle, and do so without lowering the density significantly. There is a depth-temperuture limit beyond which amphibole will not be stable, defined by its thermal-pressure stability l imit-- probably somewhat below 800°C along reasonable thermal gradients (Wyllie, 1977 ; Mysen and Boettcher, 1975a). The current depth of the 8O0°C isotherm probably is about 180 km (Johnston et al., 1974; Johnston and ToksSz, 1976).

The presence or absence of amphibole has important petrogenetic implications, because experiments have shown that first few percent partial melts of water-bearing systems can produce very silicic melts-- granitic liquids from basaltic parent mate- rial (Helz, 1976); andesite and quartz tholeiite from a lherzolite parent (Mysen and Boettcher, 1975b). Hence small amounts of relatively silicic melts may be

BASALT NOMENCLATURE AFTER BOETTCHER

Di Di

BASANITE ~ L ITNEI T E ~ S ~ O ~ ~ NEPHELINITE$ q ' ~ ~ I ~,~Q] - Thol C02 H20

Ne Oz Ne Qz

ALKALI - BASALTS FO Oliv- Thol FO

THOLEIITES

FIG. 6. Diagram of the classification of basalts (left), based on the relative abundance of CIPW (low pressure) norm minerals. Other diagram shows the composition of first partial melts of garnet peridotite produced at high pressure--dry (circle), wet (arrow to right, toward quartz tholeiite) and CO2-rich (arrow to left, toward alkali-olivine basalt).

524 McGETCHIN AND SMYTH

Fo zo75

E~RTH i ~ ~ J JMARS , ~ 4 0 kb

DAVIS 8~ SCH,~IRER 1 9 6 5

179o Di Di ,Ggo I *, °,'o) Py Di

~ " / ~ \ \ (LOW P) 2 0 k b

LeglO fo 2 = 6 KUSHIRO

Fe ' " ~c,' Fo Fo" ~ ' ' ' Qz OXIDE E.

Fro. 7. Phase diagrams relevant to partial melting of mantle compositions. The compositions of the calculated mantles for Earth (circle with cross) and Mars (circle with arrow) are shown. The first partial melts produced on heating on all three diagrams is at a ternary eutectic point, in diagram A at 1670°C, B at 1200°C, and C at 1299°C. Partial melting paths (for equilibrium) are indicated by the dotted lines. The compositions of partial melts (for 20% melting) under certain assumptions (see text) at the quaternary eutectic location are summarized in Tables VII and VIII.

produced b y small degrees of par t ia l m e l t i n g - - a n d still be consistent with the overall model we propose.

4. LAVAS PRODUCED BY PARTIAL MELT- ING OF THE MARTIAN MANTLE

The relevant experimental results on iron-rich peridoti tes do not exist, and furthermore, melt ing relations in peridoti te systems have proved to be exceedingly complex functions of pressure and water and CO2 abundance. No one has a t t empted to assess the role of sulfur in part ial melts at high pressure, which m a y be very im- por t an t in Mars. Still, much is known, at least enough to address the question of some propert ies of likely melts produced by part ial melt ing of the Mar t i an interior.

We know tha t a wide va r ie ty of melts can be produced f rom garne t -per ido t i te source material , ranging f rom si l ica-undersaturated alkali-olivine basal t (under high CO2 con- ditions, Mysen and Boettcher, 1975b; Eggler, 1975) to silica sa tura ted qua r t z -

tholeiite melts (under high H20 conditions). Green and Ringwood (1967a) have pro- duced a wide var ie ty of basalt ic melts by vary ing the degree of part ial melt ing and water pressure. Recent papers by Mysen and Holloway (1977) and Mysen and Kushiro (1977) have shown tha t 20-kbar p'~rtial melt ing of spinel-lherzolite pro- duces, progressiw~ly, a lkal~ol ivine basalt , olivine tholeiite, picrite and peridotit ic komat i i te (see Fig. 6).

Our calculated models for the mant les of both the Ea r th "rod Mars have abundan t olivine, garnet, and clinopyroxene. The iron-free ga rne t -cpx-o l sys tem ( that is, p y r o p e - d i o p s i d e - f o r s t e r i t e ) has been studied at 40 kbar (Davis and Schairer, 1965). The phase relations are simple, a te rnary eutectic sys tem (see Fig. 7A). This diagram serw~'s as a simple model for the Ea r th ' s mantle , and al though it is an iron- free system, it m a y serve as a Mars model as well if the topology of the d iagram does not alter great ly as minerals become more iron-rich.

IMPLICATIONS O F M A R S MANTLE DENSITY

TABLI~; V[

~ I ( ) D E L ~ I E L T S PI IOI ) I :CEI ) IIY PAICTI.~.L ~ { E L T I N G OF .~'I.\NTLE ASSEMttLAGES

525

Earth Mars

Parent rock Model melt, this paper

.h.[elt Parent rock Model melt, from one this paper

experiment, Kushiro

olivine 56.2 (fo 85) 5 garnet 12.7 (py 72) 32 cpx 13.7 (di 60) 33 opx 15.9 30 oxide ilm 1.3

3 73.2 (fo 67) 7 47 11.4 (py 59) 42 47 12.2 (di 49) 43

3 2.2 8 1.2

Si02 51.4 50.16 44.5 A I 2 0 ~ 8.8 11.89 11.3 MgO 24.6 25.78 17.3 FeO 4.9 - - 11.4 FelOn 2.4 - - 3.0 CaO 7.5 12.17 9.5 Na20 1.4 1.8 K 2 0 0.3 0.4

1 2 0 0 ° C 27 12 n 1027°C ]42 54

L e t us deve lop a s impl i f ied p a r t i a l m e l t - scenar io for our m o d e l m a n t l e s ( E a r t h a n d M a r s ) ; t h e E a r t h m o d e l m e l t s can be com- p a r e d w i th b a s a l t i c l a v a s as t h e y a c t u a l l y occur . I f t he r e s u l t is r easonab le , t h e n we m a y be ab le to d r a w some p l aus ib l e in- fe rence f rom the p r e d i c t e d M a r s mel t s .

F i r s t , b ecause b o t h m o d e l s cons i s t p r i n c i p a l l y of g a r n e t + o l iv ine + cl ino- p y r o x e n e ( M a r s = 97%, E a r t h 8 3 % ) , l e t us beg in w i t h t h e F o - D i - P y d i a g r a m a t 40 k b a r ( D a v i s and Scha i re r , 1965). I f t h e a d d i t i o n of i ron to t he s y s t e m does n o t a l t e r t h e d i a g r a m s ign i f i can t ly , a p l aus ib l e a s s u m p t i o n ( b a r r i n g d i s soc ia t ion of a phase ) , t h e n we can p lo t b o t h m a n t l e com- pos i t i ons on th i s d i a g r a m (see Fig . 7A). C r y s t a l l i z a t i o n a n d p a r t i a l m e l t i n g of p e r i d o t i t e c o m p o s i t i o n s are d e v e l o p e d in de ta i l in Y o d e r (1976, p. 108); for th i s d i a g r a m i t is suff ic ient to s t~ te t h a t me l t s , u p to a b o u t 20%, will be eu tec t ic , hence i sochemica l w i th t he c o m p o s i t i o n (ol7 cpx47 g46) a n d also i s o t h e r m a l .

N e x t we need to e v a l u a t e t he p r o b a b l e inf luence of t he r e m a i n i n g p h a s e (opx for the E a r t h m o d e l ; oxide for the M a r s

T °C 1500 1300 II00 I000 6 I I L I I I

i ~ _ ~%q? { . . . . . . . . . " ........... )]

\ ' I

I , I I E i J 6.0 7.0

104/ T °K

FIG. 8. Calculated viscosity as a function of tem- perature for several terrestrial lavas (after Car- michael et al., 1974); the viscosity of Martian lava (see Table VI) was calculated using the approach of Bottinga and Weill (1972).

5 2 6 M c G E T C H I N A N D S M Y T H

t2 ©

g ~

~2

v

9 o

III

H~

ffl

©

v

¢ / o g

, g r :

5"q

5"q

¢ q

zg~,i , -~ ¢'q

~ + ~ ~

. . . . . . . . . . . I I

I

" " I

I M P L I C A T I O N S OF MAP.S M A N T L E D E N S I T Y 527

I ~ 1 11 ~ 1 ~ C~

• ° °

c~

. ~

v ~

o

N ~

~'~

° ~

~ . ~ ° ~

~ m

N N ' ~ e ~ m ~ . ~ ~ ~ ~.~

~ ~ c~ " ~ . ~

~ ~" ~'~ © - ~ r',,.

~ © , . . : ,~ 21 m ~ ' ~ ~ "N~

528 McGETCHIN AND SMYTII

model). If we ignore garnet, we can examine, the Earth model in the Di-Fo-SiO2 system, studied at 20 kbar by Kushiro (1969), and the Mars model using Di-Fo-Pr , after Presnall (1966), shown in Figs. 7B and C, respectively.

The relations in the Di-Fo-SiO2 diagram are somewhat complex due to the reaction- relationship between olivine and silica-rich melts but the principal point is that first partial melts from olivine-rich bulk com- positions (such as point E in Fig. 7B, the assumed composition of the Earth mantle) occurs at the point marked 1200°C, the intersection of two cotectic lines. Cpx and opx appear to enter the melt in about equal amounts at this invariant point, until upon extended partial melting cpx disappears.

If we combine these results for the Earth, namely that the quaternary eutectic melt will consist of the eutectic ratios of the py-di-fo system (o17 cpx47 garnet4~), plus opx in approximate equal abundance to the cpx, we derive a model eutectic melt for the Earth of o15 cpx~3 opx30 g32. This composition is shown in Table VI. I t is basaltic, silica undersaturated but hyper- sthene normative, hence an olivine- tholeiite. I t is rather Al-poor and w~ry Mg-rich, and Fe-poor relative to most basalts. We have assumed no partitioning of Fe over Mg into the melt so with this in mind the result is a reasonable one; we will return to this point later. Kushiro found the 40 kbar cutectic experimentally to be (o13 epx47 opx~ g47), indicating that opx is quite reluctant to enter the melt at higher pressure; the bulk composition of the melt does not differ greatly from the one calculated here however (see Table VI).

Turning to the Mars case, we consider the Di-Fo-oxide diagram (Fig. 7C), on which we have plotted the model Mars mantle composition (M). Approximately 87% of the bulk composition of the rock is represented by the three phases olivine, cpx, and oxide; garnet is 13%. This is a low pr~ssure diagram and the phase rela-

tions are quite sensitive to f¢~; we have arbitrarily chosen Presnall's log fo2 = --6 diagram, (which is not far from the QFM buffer at these temperatures but probably substantially higher than the earth's mantle, according to Sato, 1978). The principal points are that eutectic melting does occur in the system at point A; the oxide phase is refractory and enters the melt slowly like olivine; diopside enters readily. For the bulk composition (M), posteutectic melting occurs along the cotectic boundary toward the diopside corner of the diagram, along which olivine and cpx enter the melt together. Oxide disappears before cpx for bulk composi- tion, M. In this model the oxide enters the melt slowly, more like olivine, than the cpx.

For our Mars model let us assume (1) that garnet-cpx-olivine enter the melt in proportions approximately equ'fl to the Py -Di -Fo eutectic (namely, o17 cpx47 g4~) even though the phases are iron-rich and (2) that the oxide phase enters like olivine, hence we infer a quaternary eutectie of approximately g4~ cpx43 oxides o17. This melt composition is shown in Table VI. I t is ultrabasie (SiO2 < 45%) and rich in Mg and Fe. The norm (Table VII) suggests an olivine-rich alkali basalt. It should be noted that this model is conservative in the sense that the oxide (or spinel) phase may be significantly less refractory than we have assumed here, in which case the melt would be even more silica depleted and iron-rich than this model.

There are several striking features of these postulated Martian lavas--(1) they are ultrabasic, similar in that respect to the Viking XRF soil analysis (see Table VII), (2) they would be less viscous by a factor of 2 or 3 than the model Earth lava, and a factor of 2 to 20 than terrestrial bas~flts (see Fig. S), and (3) they are rather more like komatiites than basalts (see Table VII) (Green et al., 1975; Nesbitt, 1971; Viljoen, 1969; Arndt, 1969).

Because, the inferred Martian melts are

IMPLICATIONS OF MARS MANTLE DENSITY 529

iron-rich, there is a possibility of liquid immiscibility which should be mentioned. Roedder (1951) has shown that a two- liquid field exists in the system FeO-K20, A12Oa-SiO2 ; field and petrographic evidence for such natural melts have been observed in the late stage Skaergaard granophyres of Greenland (McBirney, 1975), and in lunar samples (Roedder and Weiblen, 1970). The density contrast between the two liquids can be marked. In the Skaer- gaard ease for example, the ferrogabbro density is 2.8, the granophyre 2.4, (MeBirney, 1975). Since the melt vis- cosity is low, gravitational separation of the two melts w(mld be efficient. Martian ferro-granites could originate in this way.

It is well known that iron is strongly partitioned into magma relative to mag- nesium (Bowen and Sehairer, 1935; Green and Ringwood, 1967a; Walker et al., 1975), when crystals and liquids coexist in equi- librium. Our eMculated melts have, by definition, the same Mg/Fe ratio of the bulk parent rock. If we assume, KD ~ 0.28 (Walker et al., 1975) the terrestrial model lavas which coexist with forstcritic olivine (fo = 85) would have atom ratio (Mg/Mg 4-Fe) ~ 0.55; the Martian model lawLs, with more iron-rich olivine (fo = 67) would have (Mg/Mg + Fe) ~ 0.34. Lava com- positions having these values are shown in Table VIII. The compositions were calcu- lated using the original lava model composi- tions, but with modified (Mg/Mg 4-Fe) atom ratios; because these modified com- positions contain more iron, the other oxides are diluted on a weight basis. The result for the Earth is a more marie basalt in terms of SiO2 content--still an olivine tholeiite. This refined Mars model lava is yet more ultrabasic (SiO2 = 41.5%) and more similar to the Viking XRF soil analysis than the previous model lava. Finally, as this manuscript was being sub- mitted, Maderazzo and Huguenin (1977) published a model Mars lava composition based on their interpretation of Martian

reflection spectra, together with the Viking XRF data, which is essentially indis- tinguishable from our model (see Table VIII).

CONCLUSIONS AND DISCUSSION

(1) The apparent high density of the Martian mantle (p ~--3.55 g/cm 3) implies an iron-rich chemistry. At densities in excess of about 3.5 g/cm 3 a different phase assemblage is implied, not just higher iron within individual minerals. A large number of possibilities for the assemblages exist, depending mainly on partial pressure of w~por species (or the activity of O-S-C-H). A model composition consisting simply of pyrolite III plus 12.8 wt% FeO has the proper density (3.55 g/cm a) to match the Martian mantle; the mineral assemblage is not garnet peridotite (lherzolite), like the Earth's mantle, rather is an oxide-bearing assemblage, a rock properly called oxide- garnet, wehrlite with the following com- position-oxide (perielase-wiistite), 2%; garnet, 11% ; elinopyroxene, 12%; olivine, 73% (fo = 67), and ilmenite, 1%. All the phases are iron-rich (for KD ~- 1, Mg/Mg q- Fe = 67, atom).

Of many possible variants, perhaps the FeS-rich chemistry is the most interesting. Sulfur is enriched in the Martian soil as indicated by the Viking (XRF) inorganic chemistry experiment probably by sedi- mentary processes, howew~'r, some cosmo- logical models predict a sulfur and volatile- rich Mars. Reasonable upper limits for sulfur abundance, however, are in the 3 to 7% range, namely CC meteorite values. Sulfur reactions with olivine could yield orthopyroxene pyrrhotite-magnetite (or perhaps andradite garnet at high pressure), producing an oxide-sulfide-(garnet?)-cpx- opx and olivine-depleted assemblage (a strange websterite, see Fig. 5).

(2) Partial melting of the postulated Mars mantle (oxide garnet wehrlite) is likely to produce very basic melts--prob- ably ultrabasic in composition. A simple

530 McGETCHIN AND SMYTH

y.

O

©

~+

° ~ l Z

I

I

÷

a~ a2

® ® ® ®

IMPLICATIONS OF MARS MANTLE DENSITY 531

partial melting model based on available phase diagrams suggest that the melts will have very low silica contents (SiO2, 41 to 44%, with high iron contents) and extremely low viscosities (see Fig. 8).

(3) Liquid immiscibility is observed in Fe-rich melts. Fe-rich granophyric melts could result in this way, so volumetrically minor amounts of Martian granites could coexist with the ultrabasic melts.

(4) Some cosmological models predict a nolatile-rich Mars but it is important to vote that there is a growing body of opinion that Mars is depleted in volatiles relative to the Earth (Anders and Owen, 1977a,b,c). A volatile rich interior has several implica- tions in the context of the above. Partial melting will occur at significantly lower temperature than in a volatile-free system; furthermore, first partial melts in water- rich environments tend to be silica rich (tholeiites or even andesites), but the addition of water also tends to reduce the viscosity of these melts.

(5) A consequence of a volatile-rich Martian mantle would be the production of abundant pyroclastic volcanic products, ash--possibly of a very mafic character, ferro-kimberlite, erupted directly through the lithosphere as diatremes (see McGetchin and Ullrich, 1973). The possible importance of volcanic ash was recognized long ago by McLaughlin (see Veverka and Sagan, 1974). Kimberlite eruptions from a Mars mantle with more than the Earth's volatile content is plausible. Kimberlites on the Earth are rare and probably result from local accumulations of volatiles at the base of the lithosphere; the intrusive material varies from a carbonate-rich ultra- mafic magma to solids (comminuted mantle rock) entrained in gas, with little or no magma necessarily present. These complex volatile-rich fluids apparently are emplaced as a slurry, some with large flow speeds. They commonly are followed by more passive lava eruptions, so the source vents need not be exposed. The principal point is

that a wet Martian mantle would be a likely source for deep-seated kimberlite- like ash--perhaps in great abundance; this material could be an important constituent of the soil volumetrically.

(6) Another consequence of a wet mantle would be its effect on the thickness of the lithosphere. The lithosphere's base is prob- ably defined by the solidus, which for a dry Mars is very deep (see Johnston et al., 1974; JT, 1976). The wet solidus is a function of amphibole stability limit and could be as shallow as 100 km (for PT -- PH~o) or 200 to 250 km for 0 < Pn~o < P~. The latter estimate is in agreement with geophysical data-- the apparent finite strength of the lithosphere and partial isostatic anomaly under Tharsis (Phillips et al., 1973).

(7) According to most thermal models, melting of a dry Mars begins at about 2 By after accretion and continues to the present, gradually becoming deeper with time. In a wet Mars, the partial melt zone could ex- tend from the base of the lithosphere (cur- rently 100 to 250 km deep) downward throughout the entire mantle. Predicted temperatures are well in exccss of wet solidus temperatures; an important ques- tion is how effectively early melting events removed volatiles fromthe Martian interior.

(8) An important implication of our modeling (together with the possible wet Mars and thermal models) is that copious amounts of ultrabasic melt may have been produced within Mars. If they had access to the surface, these melts would be truly colossal lava floods.

Access of deep-seated melts to the sur- face may be tectonically controlled--func- tions of lithospheric thickness and strength, possibly its fragmentation (Hughes et al., 1977), and the stresses imposed on it by thermal expansion or contraction of the interior (Solomon and Chaiken, 1976). Currently accepted thermal models suggest that Mars is presently cooling and con- tracting, hence the lithosphere is generally

532 M c G E T C H I N AND S M Y T H

TABLE I X

MODELS FOR MINERALOGY OF UNWEATHERED MARTIAN LAVAS

Basis for model

Source of basic data used

Par t ia l melt ing of mant le mineralogy inferred from density derived from moment of inert ia

Orbital geophysics and experimental petrology

Reference McGetchin and Smyth (1977) this paper

Calculated minerah)gy

Weathering products inferred from reflection spectra and subtracted from Viking XI/F data

Planetary astronomy (reflection spectroscopy), experimental weathering of Mars sinmlants, and spacecraft da ta

Maderazzo and I iuguenin (1977)

Orthoelase 2 .4 Plagioclase 32 Olivine 37 Orthopyroxene Clinopyroxcnc 20 Nepheline 2 .5 Magnet i te 4 .4 Ihnenite trace

Rock name Pi('ritic alkali (divine Basalt,

0.6 21) 38

29 5 7

trace

Picritic alkali-olivine Basalt

in compression. Earlier in Martian history, the interior was heating up, expanding, and as a consequence the lithosphere was in tension. Because propagation of melts probably occurs by hydraulic fracturing processes driven by the buoyancy of the melt (Weertman, 1971), a tensional stress state would promote eruption to the sur- face, whereas a compressional stress state would favor sill-emplacement--that is, trapping of melt at depth with no attendant volcanism.

(9) The low viscosity of the lavas might produce higher eruption rates although flow velocities are likely to be much more sensitive to vent roughness and vent diameter (or size), than viscosity (MeGetehin and Ullrieh, 1973; Swanson, 1975). The integrated rate of production of lava over long times is clearly closely tied to the thermal history; the dynamics of an

individual eruption is probably a function of fracture or dike propagation mechanics, vent configuration, and pipe flow. The problem has two parts, the production and storage of melt at depth and the discharge during eruption.

(10) The points above have obvious implications for many surface features. One important one is that the large flood features commonly attributed to water erosion, in fact, may be largely volcanic floods, as already suggested, at least in part., by Cart (1973, 1974) and Cutts (1977).

(11) One major implication of this paper is that even though Mars occupies a place intermediate between the Earth and the Moon in most respects, such as size and mean density of the bulk planet, and many aspects of its activity (Carr, 1973, 1974; Head et al., 1977; Muteh and Saunders,

IMPLICATIONS OF MARS MANTLE DENSITY 533

1976; Masursky et al., 1974), this may not be true for volcanic processes--Mars may have a unique style of volcanism, the result of an unusual mantle (that is, an iron-rich chemistry, different mineralogy, and high volatile content).

(12) The Rasool and Le Sergeant (1977) suggestion tha t the Viking 4°A/36A ratio and argon abundance implies a low surface K20 content (of about 0.25%) is consistent with our principal conclusion, namely tha t Mart ian lavas may be ultrabasic rocks or highly picritic (olivine-rich) basalts.

(13) Finally the Mart ian lavas we pro- pose hcre (pictritic olivine alkali-basalt) are essentially indistinguishable from the model proposed by Maderazzo and Huguenin (1977), al though the models were derived completely independently and from completely different data bases (see Table IX). Both model lavas are ultrabasic in chemistry. The Madcrazzo- Huguenin model rock is based on the Viking X R F soil chemistry result modified for weathering effects (3 -5% magnetite, 5- 10% unweathered basalt, 13% brine).

(14) We would like to conclude with the remark tha t wc believe the details of the Mar t ian mantle mineralogy and partial melting processes will be strongly controlled by the role of the vapor species prcscnt and their abundance. As a consequence, the right observational data (such as Fc 3+, Fe 2+ in various mineral phases remotely sensed, or determined by in si tu observations of mineral chemistry) may constrain these models and eliminate many. In this manner, impor tant interpretations may be made which ult imately guide us in intelligently designing the s t ra tegy and tactics of eventual return of samples from Mars.

ACKNOWLEDGMENTS

This research was supported in part by NASA (Planetology Geology Programs Office and the Basaltic Volcanism Project of the Lunar and Planetary Institute) and in part by the US-ERDA (Division of Physical Research). This paper con- stitutes No. 008 of the Basaltic Volcanism Study

Project, administered by the LPI/USRA under Con- tract No. NSR 09-051-001 with the National Aeronautics and Space Administration. We profited from discussions with our colleagues, particularly in Houston and Los Alamos, and are grateful for thoughtful reviews from Robin Brett, Michael Duke, Norman Sleep, and Howard Wilshire.

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