The implications of basalt in the formation and evolution of mountains on Venus

PHYSICS O F T H E EARTH AND PLANETARY INTERIORS

ELSEVIER

Physics of the Earth and Planetary Interiors 89 (1995) 163-175

The implications of basalt in the formation and evolution of mountains on Venus Matthew G. Jull *, Jafar Arkani-Hamed Department of Earth and Planetary Science, McGill University, Montreal, Que. H3A 2A7, Canada

Received 25 January 1994; accepted 20 December 1994

Abstract The highland region of Ishtar Terra on Venus has mountains that reach up to 11 km in height and are thought to be basaltic in composition. Assuming that dynamic uplift of crust to this height is unlikely, we examine the topography produced by an isostatically supported thickening basaltic crust. It is found that regardless of whether the crust thickens by crustal shortening or by volcanic construction, the high-density basalt-eclogite phase transition is the limiting factor for producing significant elevation of the mountains. The maximum height attained by basaltic mountains depends on the nature of the basalt-eclogite phase transition. Without a phase transition, a basaltic crust must thicken to greater than 100 km to reach heights over 10 km. An instantaneous phase transition of basalt to eclogite allows a maximum topographic height of less than about 2 km. However, with a time lag of 100 Ma owing to slow rates of solid-state diffusion, our calculations show that the mountains can reach elevations greater than 10 km only if they are less than 25 Ma old. Higher temperatures within the Venusian crust may decrease the extent of the stability fields of high-density basalt phases and allow high topography if the thickening crust melts. This can occur if the radioactive element concentrations measured on the surface of Venus are uniformly distributed throughout the crust, the crust thickens to greater than 65 km, and the thickened crust is older than about 400 Ma. The conflicting results of a young age predicted for high basaltic mountains and an almost uniform surface age of 500 Ma from crater populations, coupled with similarities in bulk physical properties of Venus and Earth, suggest that the basaltic surface composition found at several landing sites on the planet may not be representative of the entire crust. We suggest that Ishtar Terra formed from the collision of continent-like highly silicic cratons over a region of mantle downwelling. Lakshmi Planum resulted from the thickening of a basaltic crust and the peripheral mountain belts formed from the collision of granitic cratons that were pulled toward a downwelling region of mantle.

1. Introduction The unimodal distribution of surface topograp h y on V e n u s is c h a r a c t e r i z e d by rolling plains,

* Corresponding author present address: Bullard Laboratories, Department of Earth Sciences, University of Cambridge, Madingley Road, Cambridge CB3 0EZ, UK.

with e l e v a t i o n s varying less t h a n a k i l o m e t r e from t h e m e d i a n p l a n e t a r y r a d i u s o f 6052 km, a n d a n u m b e r o f distinct h i g h l a n d r e g i o n s which vary in h e i g h t b e t w e e n 3 a n d 11 k m a b o v e the m e d i a n plains, b u t c o n s t i t u t e only a small f r a c t i o n o f t h e total surface a r e a o f t h e p l a n e t . T h e lack o f c h a r a c t e r i s t i c surface f e a t u r e s o f p l a t e tectonics o n V e n u s suggests t h a t the h i g h l a n d s p r o b a b l y

0031-9201/95/$09.50 © 1995 Elsevier Science B.V. All rights reserved SSDI 0031-9201(95)03015-8

164

M.G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 1(33-175

did not result from the deformation of crustal plate margins. The high correlation of topography and gravity and the large geoid anomalies over the highlands indicate an origin that is more likely to be a direct result of underlying mantle convection (Phillips, 1990; Kiefer and Hagar, 1991). This is supported by the fact that a relatively dehydrated upper mantle, which is thought to occur on Venus, would not allow for the existence of a low-viscosity asthenosphere to uncouple the lithosphere from underlying mantle circulation (Kaula, 1990; Phillips, 1990). One of the more enigmatic highlands on Venus is Ishtar Terra, which consists of a central plateau surrounded by a series of linear mountain belts. The central plateau, known as Lakshmi Planum, is 2-5 km in elevation and has a diameter of about 2000 km. It is relatively smooth and is the site of two large calderas which are surrounded by volcanic flow features. Around the perimeter of Lakshmi the mountain belts of Danu, Akna, Freyja, and Maxwell Montes rise 1.5, 3, 3, and 7 km above the central plateau. Ishtar Terra is generally regarded to represent a thickened basaltic crust (Kiefer and Hagar, 1991; Vorder Bruegge and Head, 1991; ArkaniHamed, 1993; Namiki and Solomon, 1993); however, considerable debate remains over the timing and mechanism of its formation. A region of downwelling mantle has been suggested as a means for producing Ishtar (Binschadler and Parmentier, 1990; Kiefer and Hagar, 1991), whereby radial crustal shortening has caused the plateau and surrounding mountain belts to form from converging crust. A region of upwelling mantle has also been suggested as a mechanism for formation (Basilevsky, 1986; Pronin, 1986), whereby the mountain belts have formed through crustal thickening caused by shear along the base of the lithosphere from material flowing away from the upwelling. The presence of parallel ridges adjacent to the mountain belts in Lakshmi Planum and a lack of radial or concentric faulting patterns tend to suggest that the mountains and central plateau of Ishtar resulted from a compressional origin and not from crustal flow away from a region of mantle upwelling (Kaula et al., 1992).

It is doubtful that underlying mantle flow provides any appreciable dynamic support for the mountains of Ishtar. This is because the most abrupt changes in elevation, such as at Maxwell Montes, occur over distances of tens to hundreds of kilometres and probably do not reflect scales of mantle processes (Smrekar and Solomon, 1992). The correlation of gravity and topography over the much longer wavelength of Lakshmi indicates that there may be some dynamic support of the central plateau; however, this would be limited to about 4 km in elevation (Grimm and Phillips, 1991). Vorder Bruegge and Head (1991), ArkaniHamed (1993) and Namiki and Solomon (1993) examined the height of mountains resulting from an isostatically supported, thickening basaltic crust. They found that a basaltic composition severely limits the height attained by the mountains. This is because basalt, with a density of 2900-3100 kg m 3 is unstable at 10-15 kbar, and transforms through a garnet granulite mineral assemblage to eclogite, with densities of 34003500 kg m 3 at pressures greater than 15-30 kbar (Ahrens and Schubert, 1975). Therefore the thickening basaltic crust becomes considerably heavier as its roots enter the stability fields of the higher-density basalt phases. Vorder Bruegge and Head (1991) and Namiki and Solomon (1993) found that disequilibrium of basalt and eclogite is the only way that mountains of 11 km can be produced. This occurred if the mountains were less than about 50 Ma old, implying that vigorous mantle convection and active tectonism is recent in the history of Venus. However, Magellan images have revealed a nearly uniform distribution of craters on Venus, showing only minor tectonic activity in the last 500 Ma (Schaber et al., 1992). This suggests that a 50 Ma age for basalt mountains may not be a feasible explanation for their great height. In this paper we closely examine the factors which affect the height of mountains formed from a thickening basaltic crust. Three different crustal thickening models are considered. Thrust faulting of a brittle crust is modelled by stacking basaltic layers onto the surface; shortening of a ductile crust is modelled by uniformly straining the en-

M.G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175

tire crust, and volcanic construction, such as from shield volcanism, is modelled by stacking molten basaltic layers onto the crust. The temperature and density profiles are calculated within the thickening crust and are used to determine the resulting topography. A maximum height of about 2 km is achieved when the crust is 35 km thick if basalt transforms to eclogite instantaneously, as was also concluded by Arkani-Hamed (1993). Higher heat generation rates in the basaltic crust result in partial melting in the deeper parts of the crust as well as higher topography, and can account for volcanic activity at the surface if the crust is thicker than about 65 km and is older than about 400 Ma. Sluggish reaction rates for the transition of basalt to eclogite allow shortlived topography that is higher than 10 km if the crust thickens to greater than 100 km. The difficulty in maintaining high basaltic mountains for anything more than about 25 Ma in our models, and the constraints on surface age from cratering densities, suggests that Maxwell Montes are probably not formed entirely from basalt, but are more probably silicic in composition.

2. Crustal thickening To examine the effects of basaltic crustal thickening on surface topography, we model the time-dependent temperature and density profile within a thickening crust that overlies a depleted mantle layer. The temperature profile is calculated by solving the one-dimensional heat conduction equation in a Lagrangian co-ordinate system using a finite difference method. At each time step, we solve for the temperature profile within the crust, thicken the crust, and then stretch the temperature profile to match the new thickness of the crust. The temperature profile is calculated from the surface to 200 km depth, which is assumed to be the base of the thermal boundary layer of mantle convection (Schubert et al., 1990; Leitch and Yuen, 1991). The upper boundary is held at a constant temperature of 750 K, and the lower boundary is held constant at either 1600 K or 1900 K. The partial melting of

165

ascending mantle at these temperatures produces a basaltic crust of 5 km and 20 km thickness, respectively, and an associated depleted mantle residue of 40 km and 90 km thickness, respectively (Arkani-Hamed, 1993). In our models we assume thermal expansivity for the crust and mantle to be 3.0 X 10 -5 K -1, a constant specific heat capacity of 1200 kJ kg -1, and a latent heat of fusion for basalt of 400 kJ kg -1. The thermal conductivity is assumed temperature dependent, and the thermal conductivities of basalt and olivine (Touloukian and Ho, 1981) are adopted for the crust and mantle. The calculated crustal heat generation rate of 1.25 x 10-10 W kg-l, based on the abundance of radiogenic elements on the surface of Venus (Surkov et al., 1987), is significantly larger than the 2.63 x 10 -11 W kg -1 for terrestrial tholeittic basalts, and is more comparable with 9.60 x 10 -10 W kg -1 for granite. We therefore compare the effect of heat generation rates of 1.25 x 10-1° W kg -1 and 2.63 x 10 -11 W kg -1. The heat source is assumed to be uniformly distributed within the crust, depleted mantle, and undepleted mantle, with relative concentrations of radioactive isotopes in the depleted and undepleted mantle given by Turcotte and Schubert (1982, p. 141) for oceanic lithosphere. As the temperature profile and total thickness of the crust change, we recalculate the density profile of a crustal column to determine topography based on Airy isostatic support. The depth of compensation is taken to be constant at 120 kin, but a 200 km compensation depth is also considered. We consider thickening of the crust through both volcanic construction and crustal shortening. Volcanic construction is modelled by adding layers of molten rock to the surface of the crust (Fig. l(a)). Crustal shortening is modelled with both a thrust faulting model, whereby cold layers of rock are added to the surface (Fig. l(b)) and uniform thickening model, whereby the crust strains uniformly throughout (Fig. l(c)). The thrust faulting model is identical to the volcanic construction model except that the layers added to the surface are cold and contain no heat of fusion. The thrust faulting and volcanic models are referred to as the hot and cold layering models. In the uniform

166

M.G..lull, J. Arkani-Hamed /Physics of the Earth and Ptanetao, Interiors 89 (1995) 103-175

strain model we use strain rate between 1 x 10-14 s-~ to 1 x 10-~6 s-~ throughout the crust, and in the layering models the thickening rates vary from 0.5 to 5.0 km Ma-~. The basalt-granulite-eclogite phase relationships are based on the results of lto and Kennedy (1971) for terrestrial basalts. The density increase through the transition of basalt to eclogite is based on Green's (1967)

experimental results. The densities, taken at room temperature and pressure, are 2900 kg m -3 for basalt, 3250 kg m - 3 for peridotite, 3250 kg m- 3 for depleted peridotite, and 3500 kg m -3 for eclogite. All models begin to thicken after an initial period of 50 Ma where the temperatures are in equilibrium and time starts with the onset of crustal thickening.

8

b

C

Fig. I. Schematic illustrations of crustal thickening models. The crust thickens by volcanicconstruction (hot layering model) in (a) thrust faulting of a brittle crust (cold layering model) in (b), and uniform strain in (c).

M.G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175

E

20.0

,

i.-ri..u ...¢. o

"1[2. <

O [2. O

10.0

0.0

0.0 200.0 TIME (Ma)

Fig. 2. Topography from a cold crustal layering model thickening to 50, 100, and 150 km in 100 Ma with a crustal heat source of 2 . 6 3 x 1 0 -1I W kg -1, and without high-density basalt-eclogite transitions. T h e abrupt change in the curves at 100 Ma occurs because the crust stops thickening. T h e slight rise in topography after 100 Ma is from temperature increases in the thickened crust.

3. Numerical results

3.1. The basalt-eclogite transformation A thickening basaltic crust that does not undergo high-density phase changes can reach very high elevations. Fig. 2 shows topography produced by a basaltic crust thickening to 50, 100, and 150 km with a heat source of 2.63 x 10 -~1 W kg -~. After the thickening stops at 100 Ma, the topographic height remains relatively constant; it increases only slightly as a result of temperatures re-equilibrating within the crust. The maximum height attained is dependent only on the final crustal thickness. The crust must be greater than 100 km thick for topography to be higher than 10 kin. Fig. 3 shows the relationship of topographic height with composition and temperature in the underlying crust and upper mantle for a thickening model that incorporates instantaneous basalt-eclogite phase transitions. The model calculations are 1-D and so the graphs show the evolution through time of topography, composition, and temperature. After the crust begins to thicken, the topography increases in height (Fig. 3(a)) but then quickly rolls over and begins to

167

decrease as the denser granulite (2) and eclogite (3) phases replace basalt (1) in the lower parts of the crust (Fig. 3(b)). When the crust has stopped thickening after 100 Ma, eclogite has formed at the base of the crust and the isotherms are not in equilibrium (Fig. 3(c)). As temperatures re-equilibrate and the crust heats up, the amount of eclogite decreases and the topography rebounds upward. The effects on topography of thickening time, compensation depth, initial and final thickness, and heat source are summarized in Fig. 4. The maximum topographic height attained in each model is represented during thickening by a black bar and after 600 Ma by a dark grey bar. The light grey bars give the corresponding topographic height attained if there are no high-density basalt phase changes. In the cold layering model the rate of thickening does not have a significant effect on the maximum topographic height. The maximum height during formation varies from 1.3 to 1.5 km, corresponding to thickening rates of 6 km Ma-1 and 0.4 km Ma-1. The hot layering model shows higher topography (2.0-4.0 km) if the thickening rate is fast (2.0-6.0 km Ma -~) because the crust is heated by the molten layers being added to the surface, which allows less of the denser basalt phases to form while contributing to greater thermal buoyancy. Thickening rates slower than about 1.5 km Ma-1 allow the heat to be radiated out into the atmosphere faster than it can be transferred to the crust. The uniform strain model produces topography nearly identical to that for the cold pile-up model. For all models, the topographic height at 600 Ma is strongly dependent on the final thickness of the crust. For a heat source of 2.63 × 10-11 W kg-1, basaltic crusts that thicken past 60 km are negatively buoyant because of the increased amount of eclogite at their base. The optimum crustal thickness for maximum topographic height in the models with a heat source of 2.63 x 10 -11 W kg -1 is about 35 kin, and occurs when the mean density contrast between the crust and mantle is highest. Thicker initial crusts cause lower topography during formation because the base of the crust reaches the eclogite stability field sooner and because the

168

M.G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175

heat source of 2.63 x 10- f~ W kg- '. During formation the topography maximum is nearly identical for both heat sources; however, after formation temperatures can increase to such an extent that partial melting occurs within the crust. This is observed when the crust thickens beyond about 65 km and is older than about 400 Ma. Three models are considered for the history of the crust after melting and are shown by the splitting of the topography curve into three branches in Figs. 5(b) and 5(c). Complete containment of buoyant melt within the root of the thickened crust (mc in Fig. 5) causes the topography to reach elevations higher than 10 kin. Lateral extrusion of melt (rex in Fig. 5) removes latent heat of melting, thins the crust, and levels the topography at about 7 km. Vertical extrusion of melt onto the surface

pressure at a given depth below the lithosphere is lower than for a thinner initial crust. An initial crustal thickness of 20 km gives topography about 2 km lower than if the crust were initially 5 km thick.

3.2. Effect of heat source If the heat source in the crust of Venus is taken to be 1.25 × 10-~° W kg-~, based on surface measurements from Venera and Vega missions (Surkov et al., 1984, 1987), a significantly different temperature distribution and topography could be expected. Fig. 5 shows thickening to 50, 100, and 150 km in a time of 100 Ma with a heat source of 1.25 X 10-~0 W kg-~. The dashed curve shows the topography corresponding to a

50.

a I--

2: (D

00,

Iii ::E

-5o i 0.0

b 100.0

E T In Iii

200.0 0.0

]-~,:!

C

aoo .000 ¸

1000 I~oo

~200

200 0

. 0.0

.

.

.

. . 200.0

400 0

600 0

T I M E (Ma)

Fig. 3. Comparison of topographic height (a), lithosphere and upper-mantle composition (b), and temperature distribution (c) for a thickening basaltic crust that undergoes instantaneous high-density phase transitions with a heat source of 2.63 × 10--11 W kg ~. The topography in (a) initially increases with crustal thickening but then rolls over and begins to decrease with the formation of granulite and eclogite in the lower crust. In (b) the layers shown are basalt (1), granulite (2), eclogite (3), depleted mantle (4), and undepleted mantle (5). (c) shows isotherms in degrees Kelvin to a depth of 200 km. The bold lines in (c) refer to the upper and lower boundary of the undepleted mantle region.

M.G. ,lull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175 (mv in Fig. 5) also removes latent heat of melting, but maintains a constant crustal thickness. The topography is slightly lower than for lateral extrusion because the crust remains thick and because lower temperatures are pushed downwards from the layering of basalt onto the surface. The phase stability regions for the three crustal melting models are shown in Fig. 6 for 100 Ma thickening to 100 km. In all three cases the crust thickens and granulite (2) and eclogite (3) initially form at the base of the crust. As temperatures rise, eclogite decreases and basalt (1) increases in extent. After about 400 Ma, melting (6) first occurs in the deeper parts of the crust. Complete containment of melt (me) is shown in Fig. 6(a), vertical extrusion (mv) is shown in Fig. 6(b), and lateral extrusion (mx) with crustal thinning is shown in Fig. 6(c). The lowering of crustal temperatures from surface layering of extruded melt in Fig. 6(b) is indicated by the decrease in basalt phase at about 400 Ma; after this time, continual resurfacing maintains a constant temperature and

,o :

169

phase distribution. The effects of a heat source of 1.25 x 10 -]o W kg -1 on maximum topographic height is summarized in Fig. 4 for one of the thickening models. Only with complete containment of melt within the crust as it thickens to greater than 100 km can the topography reach elevations higher than 10 km. 3.3. Effect o f basalt-eclogite transformation time lag Solid-state diffusion has been examined as a possible reaction mechanism for the basalt-eclogite transition (Ahrens and Schubert, 1975). Reaction times are found to be around 10-100 Ma for temperatures above 900 K. We incorporate this effect into our models by assuming a time lag of 100 Ma over which the density of basalt increases linearly to the density of eclogite. The density change depends only on the accumulated time spent at the pressure and temperature of the stable phase. Fig. 7 shows the effect of the

:

i'OR:)GRAPHIC HEIG.HTABOVE MEAN PLANETARY SURFACE(km)

J,,,:,, J .5-

THICKENING TIME[Mg) FINALCRUSTAL

251100125oo125oc 50 'j,

I 1so

25 Im0125 I100 25 I10C I'--I 1" I

s( I

100

5ol @so

INITIAl.CRUSTAL THICKNESS(kin)

5

5

5

~PENSATION DEPTH(kin)

120

120

120

CRUSTALn.IICKENING MODEL ¢lW~t,~l. HEM

, ~ , ¢ ~ ~,'k~)

COLD LAYERING

HOT LAYERING

UNIFORM STRAIN

2.63x 10""

100'100

i

100

I oCso 5o1 oo1 5o 5

20

200

120 COLD LAYERING

5

120

....

COLD LAYERING

1.25x 10"~°

Fig. 4. Maximum topographic height during thickening (represented by black bars) and after 600 Ma (represented by dark grey bars) for a suite of crustal thickening models with an instantaneous basalt-eclogite transition. Light grey bars show corresponding topographywithout high-densityphase transitions. The letters mc, mx, and mv for a heat source of 1.25 × 10-10 W kg- ~ represent different models of crustal melting (see Figs. 5 and 6 for explanation).

170

M.G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetat3, Interiors 89 (1995) 163-175

time lag on topography for crustal thickening to 100 km in a time of 25 Ma (Fig. 7(a)), 50 Ma (Fig. 7(b)), and 100 Ma (Fig. 7(c)) with a lateral melt extrusion model. A time lag of 100 Ma allows topography to reach 10 km only for a very short period of time if the crust thickens in 25 Ma. The increase and levelling off of topography after 400 Ma is caused by the temperature increase in the crust and subsequent extrusion of melt and thinning of the crust. A summary of the effects of a time lag of 100 Ma are shown in Fig. 8 for one of the thickening models. The crust must thicken at

a rate greater than 4 km Ma greater than 100 km to attain than 10 kin. High topography time lag of 100 Ma will last not 25 Ma.

I to a thickness elevations higher produced with a longer than about

4. Discussion and conclusions We have examined the topography produced by a thickening basaltic crust with numerical models that incorporate the effects of style and

5.0

ta

0.0

-5.0 t::

15.0 • -.

II

(_9 uJ -F 0 "113._ ,< n.-

b

10.0 mx

JJ

5.0

,

mv

o.o

0

0

I--

i

-5.o

J

15.0

//

e 10.0

i

/

_

- -

mx

. . . . . . . .

mv

. . . .

5.0 /

/

----

i

0.0

!

// -5.0

/

0.0

i

20(3.0

400.0

600.0

TIME (Ma)

Fig. 5. Topographic height produced by a thickening basaltic crust with instantaneous high-density phase changes and a heat source of 1.25 × 10 - l ° W kg -1. Models shown are for thickening to 50 km (a), 100 km (b), and 150 km (c) in 100 Ma. In (b) and (c), the branches of the curve show three models of crustal melting. Complete melt retention at the base of the crust is indicated by mc, lateral extrusion and thinning of the crust is indicated by mx= and vertical extrusion of melt onto the surface is indicated by my. The dashed lines show topography with a basalt heat source of 2.63 × 10 -11 W kg ~.

M.G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175

rate of thickening, initial and final crustal thickness, depth of isostatic compensation, the basalt-eclogite phase transition, crustal radiogenic heat source concentrations, and the consequences of crustal remelting. Without the basalt-eclogite transition, topography higher than 10 km can be produced as long as the crust thickens to greater than about 100 km. Instantaneous basalt-eclogite transitions lower the maxim u m topographic height to less than 2 km when the crust is about 35 km thick, whereas a crust thicker than 35 km can produce negative topography. These results are in good agreement with those found by A r k a n i - H a m e d (1993). Higher temperatures resulting from hot crustal

171

layering allow topography up to 4 km in height only for thickening rates greater than about 4 km M a - 1 , although most of the heat is radiated into the atmosphere. Cold layering and uniform strain models do not produce appreciable differences in topographic height. With a heat source of 1.25 × 10-t0 W kg-1, remelting at the base of the crust can occur after about 400 Ma from the onset of thickening if the crust thickens to greater than 65 km. Maintaining buoyant melt within the crust allows topography higher than 10 km; however, it is unlikely that the melt could be contained for any appreciable length of time. Lateral extraction of melt causes lower topography and thins the crust. Vertical extrusion of the melt onto the

0.0

a 100.0

200.0

E ,¢,

0.0

b

V

I

100.0

O_ W

0

200.0

0.0

C 100.0

200.0 0.0

100.0

200.0

300.0

400.0

500.0

600.0

700.0

TIME (Ma) Fig. 6. Phase diagrams corresponding to the three different melt models used in Figs. 5(b) and (c). The layers are basalt (1), granulite (2), eclogite (3), depleted mantle (4), undepleted mantle (5), and partial melt (6). (a) shows complete retention of melt (mc), (b) shows vertical extrusion of melt onto the surface (mv), and (c) shows lateral extraction of melt (mx) and associated thinning of the crust to 65 km.

M. G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175

172

~v I'-r-

(.9

15.0 10.0

I..U "1-

(2_

5.0

rr (.9 0

0.0

0

-5,0

'1" 13. ,<

o_

0.0

200.0 TIME (Ma)

400.0

Fig. 7. Topographic height produced for a thickening basaltic crust with a 100 Ma time lag for the transition of basalt to eclogite. Models shown are for thickening to 100 km in 25 Ma (Curve a), 50 Ma (Curve b), and 100 Ma (Curve c) with lateral extraction of melt. The rise in topography after 400 Ma is from increased temperatures, and the flattening of the curve at about 400 Ma is from the lateral extrusion of melt.

surface maintains crustal thickness but causes even lower topography than for lateral extrusion because lower temperatures are pushed down.d

10 * TOPOGRAPHIC HEIGHT ABOVE MEAN PLANETARY SURFACE (kin)

5

"

o-

I

-5--

THICKENING TIME (Ma) FINAL CRUSTAL THICKNESS (km) INmAL CRUSTAL THICKNESS (kin) COMI~NSATION DEPTH (km)

CRUSTAL THICKENIN~ MODEL CRUSTAL HEAT SOURCE ~//kg)

5 120 HOT

LAYI~RING 1.25x10 m

Fig. 8. Maximum topographic height during thickening (represented by black bars) and after 600 Ma (represented by dark grey bars) for a crustal thickening model with a 100 Ma time lag for the basalt-eclogite transition and a heat source of 1.25×10 -1° W kg -t.

wards. Incorporating a time lag of 100 Ma into the basalt-eclogite transition allows topography higher than 10 km for less than 25 Ma if the crust thickens to greater than 100 km. Our conclusions about the effect of finite reaction rates for the basalt-eclogite transition on topography do not differ significantly from the results of Namiki and Solomon (1993), who predicted a young age of 50 Ma for Maxwell Montes, based on modelling the basalt-eclogite transition with an equation for solid-state diffusion. Their results depend on estimation of cation diffusion coefficients and grain size in the basalt, giving slightly longer periods of disequilibrium than with our linear time lag. The minimum formation time of 25 Ma in our models clearly indicates, as does Namiki and Solomon's result of 50 Ma, that phase disequilibrium allows high basaltic mountains only for very short periods of time. Such a result is in direct conflict with cratering records of the surface, which indicate that the crust has probably not deformed significantly within the past 500 Ma (Schaber et al., 1992). Despite the fact that Maxwell Montes cover a small region of the planetary surface, it is doubtful that they could escape dating of 500 Ma because the amount of crustal material required to produce such a large massif would be considerable. The extent of this region, assuming that the crust has thickened to 100 km under Maxwell Montes, and that its areal extent is 900 km × 500 krn, would be about 4.5 × 10 ~ km 2 if undeformed crust is about 10 km thick. This suggests that a region more than 10 times the area of these mountains was involved in their formation, if they were formed by crustal thickening. This clearly indicates that it is difficult for Maxwell Montes to avoid an age of 500 Ma from global cratering averages. Also conflicting with recent tectonic activity is the presence of the impact crater Cleopatra in Maxwell Montes. Cleopatra is 100 km in diameter and is undeformed in appearance; this suggests that either it is very recent, in which case it does not constrain the age of Maxwell, or that it is old and the surrounding crust has not been substantially deformed since its formation. The improbability (only about 8% (Namiki and Solomon, 1993)) that the crater and Maxwell are both very young,

M.G. Jul~ J. Arkani-Hamed / Physics of the Earth and Planetary Interiors 89 (1995) 163-175

combined with implication of involving a large area of crust to form Maxwell, supports an older age for the mountains. Abundant evidence of fresh-looking volcanic flows in the plains of Lakshmi Planum (Kaula et al., 1992) can also help to put a constraint on the age of Ishtar. The calderas Colette and Sacajawea in Lakshmi Planum appear amidst smooth, undeformed volcanic plains which embay ridged terrain. Assuming that Ishtar is situated over a site of mantle downwelling, the lower temperatures of the downwelling mantle would not be expected to melt the lower crust; it is more likely that the crust melted from its own heat source. The high latent heat of 850 J kg-1 K-1 and low mantle temperatures of 1300-1400 K in the models of Namiki and Solomon (1993) do not allow crustal melting to occur. However, the values used in our models of 400 J kg- 1 K - 1 and 1600 K for latent heat and mantle temperature allow for melting of the crust after about 400 Ma. The apparently recent volcanic flows in Lakshmi may therefore be the result of crustal thickening that occurred long ago and only recently have crustal temperatures risen sufficiently to cause melting of the lower crust. This result also explains how volcanic activity can occur in an old and tectonically inactive area. The apparently irreconcilable result of an age significantly greater than 50 Ma for Maxwell Montes with a basaltic composition that undergoes high-density phase changes may be avoided if they do not consist of basalt. The assumption that Maxwell Montes are formed from basalt is based on surface analyses taken at three different locations on the planet which showed similar compositions to terrestrial basalts. On Earth, however, approximately 30% of the crust is composed of permanently buoyant, silicic continental crust. Considering that Venus and Earth are likely to have very similar bulk physical properties, it seems a gross simplification to assume that the mantle of Venus has not been able to differentiate sufficiently to produce more evolved silicic crustal components. The production of silicic continental crust on Earth is a widely controversial subject. It is generally believed that subduction of oceanic litho-

173

sphere plays a key role in the formation of silicic magmas by introducing water into the upper mantle (see Gill, 1981, pp. 116-121 for a summary). Under hydrous conditions, crystallization fractionation of a basalt parent magma causes silica and alkali enrichment and iron depletion (calc-alkaline magma series), whereas under anhydrous conditions basalt magmas differentiate with iron enrichment and only modest increases in silica content (tholeiitic magma series) (Sisson and Grove, 1993). Support for the role of subduction in the generation of silicic crust is provide by the abundance of rhyolite volcanism at island arcs and the presence of large granitic batholiths at subduction zones along continental margins, such as the Sierra Nevada and Andes mountains of North and South America. In comparison, at spreading ridges, where melting is caused by decompression melting of the upper mantle, there is little evidence of rhyolitic volcanism. The similarities in trace element depletion of Ti, Ta, and Nb in continental cratons with those of silicic magmas produced at island arcs gives strong evidence that subduction of oceanic lithosphere plays a key role as a source for the silicic crustal material of the continents. In light of the requirements needed to produce silicic magmas on Earth, the absence of plate tectonics on Venus and the relative dearth of water in its atmosphere tends to suggest an environment in which silicic magmas are relatively scarce. Early in its history, however, Venus may once have had enough water to form a terrestrial-style ocean (Donahue et al., 1982; Donahue and Hodges, 1992). Based on measurements of the ratio of deuterium to hydrogen in the atmosphere, Donahue et a1.(1982) concluded that as a lower bound the atmosphere of Venus once contained two orders of magnitude more water than it does today. This enrichment of H 2 would be equivalent to about 0.3% of a terrestrial ocean (Donahue et al., 1982). Therefore, very early in its history Venus may have been more similar to Earth than it is now. At a time of reduced solar luminosity, water may have existed on the surface of the planet, possibly to the extent of a terrestrial-style ocean. If plate tectonics were once active on Venus, silicic crustal

174

M. G. Jull, J. Arkani-Hamed / Physics of the Earth and Planetary" Interiors 89 (1995) 163-175

material may have been produced by the subduction of hydrated crust, similar to how it is produced on Earth. However, even if plate tectonics never operated on Venus and there was no transport of water into the upper mantle through subduction, remelting of a thickened, hydrated crust and fractionation of primary magmas from the mantle could have allowed for increased silica content of melt which was erupted onto the surface. Experiments have also shown that under dehydrated conditions partial melts produced from peridotite can become silica enriched through the fractionation of F e - T i oxides (Johnston, 1986). Support for silicic volcanism on Venus is provided by observations of pancake-like lava domes, the morphology of which has been attributed to the result of a silicic composition (e.g. Fink et al., 1993). On Venus, silicic crust could therefore have formed, possibly congregating into small cratons. At later times the vaporization of water on Venus may have hindered the formation of silicic crust and as a consequence fewer continental-type cra-

tons were formed. As mantle convection became less turbulent at around 1-0.5 Ga ago and began to settle into a steady state, the isolated small cratons may have independently converged above a region of mantle downwelling. Their collision would have resulted in the formation of the peripheral mountain belts of Ishtar, with a thickened basaltic crust forming the central plateau of Lakshmi. The variation in distance of the mountains from the downgoing central part may have resulted from the different arrival times of the various cratons. Maxwell Montes could then have formed and maintained a high elevation because their silicic composition would not have undergone high-density phase changes, and remelting at the base of the thickened basaltic plateau after a period of about 400 Ma could account for the fresh-looking volcanic features on Lakshmi Planum. Fig. 9 shows a schematic depiction of the type of scenario that could have formed a highland such as lshtar. This scenario reconciles the conflicting observations of high elevations and old ages for the mountains of Ishtar, and the

Fig. 9. Schematic depiction of the formation of a highland such as Ishtar Terra from the convergence of granitic cratons over a region of mantle downwelling. The high peripheral mountain belts are formed from buoyant, silicic crust. Volcanism in Lakshmi Planum is the result of melting near the base of a thickened basaltic crust.

M.G. Jull, J. Arkani-Hamed /Physics of the Earth and Planetary Interiors 89 (1995) 163-175

relatively flat plateau and fresh-looking volcanic flows of Lakshmi, and helps to identify a compositional variation in the lithosphere of Venus.

Acknowledgements This research was supported by the Natural Sciences and Engineering Research Council (NSERC) of Canada, under Operating Grant OGP0041245, and by a grant to M.G. Jull from the Department of Earth and Planetary Sciences at McGill University.

References Ahrens, T.J. and Schubert, G., 1975. Gabhro-eclogite reaction rate and its geophysical significance. Rev. Geophys. Space Phys., 13: 383-400. Arkani-Hamed, J., 1993. On the tectonics of Venus. Phys. Earth Planet. Inter., 76: 75-96. Basilevsky, A.T., 1986. Structure of central and eastern areas of Ishtar Terra and some problems of Venusian tectonics. Geotectonics, 20: 282-288. Binschadler, D.L and Parmentier, E.M., 1990. Mantle flow tectonics: the influence of a ductile lower crust and implications for the formation of topographic uplands on Venus. J. Geophys. Res., 95: 21329-21344. Donahue, T.M. and Hodges, Jr., R.R., 1992. Past and present water budget of Venus. J. Geophys. Res., 97: 6083-6091. Donahue, T.M., Hoffman, J.H., Hodges, Jr., R.R. and Watson, A.J., 1982. Venus was wet: a measurement of the ratio of deuterium to hydrogen. Science, 216: 630-633. Fink, J.H., Bridges, N.T. and Grimm, R.E., 1993. Shapes of Venusian 'pancake' domes imply episodic emplacement and silicic composition. Geophys. Res. Lett., 20: 261-264. Gill, J.B., 1981. Orogenic Andesites and Plate Tectonics. Springer, Berlin. Green, D.H., 1967. Effects of high pressure on basaltic rocks. In: H.H. Hess and A. Poldervaart (Editors), Basalts: the Poldervaart Treatise on Rocks of Basaltic Composition, Vol. 1. Wiley-Interscience, New York, pp. 401-443. Grimm, R.E. and Phillips, R.J., 199l. Gravity anomalies, compensation mechanisms, and the geodynamics of western lshtar Terra, Venus. J. Geophys. Res., 96: 8305-8324. Ito, K. and Kennedy, G.C., 1971. An experimental study of the basalt-garnet granulite-eclogite transition in the structure and physical properties of the Earth's crust. Geophys. Monogr. Am. Geophys. Union, 14: 303-314. Johnston, A.D., 1986. Anhydrous P - T phase relations of

175

near-primary high alumina basalt from the South Sandwich Islands. Contrib. Mineral. Petrol., 92: 368-382. Kaula, W.M., 1990. Mantle convection and crustal evolution on Venus. Geophys. Res. Lett., 17: 1401-1403. Kaula, W.M., Binschadler, D.L., Grimm, R.E., Hansen, V.L., Roberts, K.M. and Smrekar, S., 1992. Styles of deformation in Ishtar Terra and their implications. J. Geophys. Res., 97: 16085-16120. Kiefer, W.S. and Hagar, B.H., 1991. Mantle downwelling and crustal convergence: a model for Ishtar Terra, Venus. J. Geophys. Res., 96: 20967-20980. Leitch, A.M. and Yuen, D.A,, 1991. Compressible convection in a viscous Venusian mantle. J. Geophys. Res,, 96: 15551-15562. Namiki, N. and Solomon, S.C., 1993. The gabbro-eclogite phase transition and the elevation of mountain belts on Venus. J. Geophys. Res., 98: 15025-15031. Phillips, R.J., 1990. Convection-driven tectonics of Venus. J. Geophys. Res., 95: 1301-1316. Pronin, A.A., 1986. The structure of Lakshmi Planum, an indication of horizontal asthenospheric flow on Venus. Geotectonics, 20: 271-281. Schaber, G.G., Strom, R.G., Moore, H.J., Soderblom, L.A., Kirk, R.L., Chadwick, D.J., Dawson, D.D., Gaddis, J.M., Boyce, J.M. and Russell, J,, 1992. Geology and distribution of impact craters on Venus: what are they telling us? J. Geophys. Res., 97: 13257-13301. Schubert, G., Bercovici, D. and Glatzmaier, G.A., 1990. Mantle dynamics in Mars and Venus. J. Geophys, Res., 95: 14105-14129. Sisson, T.W. and Grove, T.L., 1993. Experimental investigations of the role of H20 in calc-alkaline differentiation and subduction zone magmatism. Contrib. Mineral. Petrol., 113: 143-166. Smrekar, S. and Solomon, S.C., 1992. Gravitational spreading of high terrain in Ishtar Terra, Venus. J. Geophys. Res., 97: 16121-16148. Surkov, Y.A., Barsukov, V.L., Moskalyeva, L.P., Kharyukova, V.P. and Kemurdzhian, A.L., 1984. New data on the composition, structure, and properties of Venus rock obtained by Venera 13 and Venera 14. Proc. 14th Lunar Planet. Sci. Conf., Part 2. J. Geophys. Res., 89: B393-B402. Surkov, Y.A., Kirnozov, F.F., Glazov, V.N., Duchenko, A.G., Tatsy, L.P. and Sobornov, O.P., 1987. Uranium, thorium, and potassium in the Venusian rocks at the landing sites of Vega and 2. Proc. 17th Lunar Planet. Sci. Conf., Part 2. J. Geophys. Res., 92: E537-E540. Touloukian, Y.S. and Ho, C.Y., 1981. Physical Properties of Rocks and Minerals. McGraw-Hill, New York. Turcotte, D.L. and Schubert, G., 1982. Geodynamics: Applications of Continuum Physics to Geological Problems. Wiley, New York. Vorder Bruegge, R.W. and Head, J.W., 1991. Processes of formation and evolution of mountain belts on Venus. Geology, 19: 885-888.