Geological Evolution of Venus:Rises, Plains, Plumes, and Plateaus

Roger J. Phillips* and Vicki L. Hansen

Crustal plateaus and volcanic rises, major physiographic features on Venus, both formedover mantle plumes. Crustal plateaus were produced by large degrees of plume meltingbeneath thin lithosphere. The oldest tectonic features in crustal plateaus are ribbon-liketroughs indicating early uplift and tensile stretching; their shallow depths suggest thatsurface temperature there was higher in the past. Widespread volcanic plains, derivedfrom the broad upwellings of internally heated mantle convection, were continuallyerupted during the time of crustal plateau formation. Discrete volcanic rises, younger thancrustal plateaus, formed over a thicker lithosphere, with less plume partial melting. Thepresence of only one transitional form indicates that the lithosphere thickened rapidly.Thermal and magmatic models show that the formation of these major features can betied to an expected change in mantle convective style about 1 billion years ago.

The recognition of the significance ofjust two major terrestrial tectonic featuresin the oceans—midocean ridges andtrenches—established the paradigm of sea-floor spreading and subduction. This recog-nition in turn provided the kinematic foun-dation for plate tectonics, an elegant theoryproviding a coherent framework for under-standing how Earth’s broad geomorphic fea-tures—including volcanic chains and con-tinental mountain belts—relate to heat es-cape and interior processes. Similarly, un-derstanding the differences in the formationof two major features on Venus—crustalplateaus and volcanic rises—is key for un-derstanding the geological evolution of thatplanet over the last billion years. In thisarticle, we examine the geology of thesefeatures and present a model for theirformation and of the thermal evolution ofthe lithosphere (1) of Venus. We suggestthat the features both formed as a result ofthe interaction with the lithosphere ofmantle plumes rising from the core-mantleboundary.

Volcanic plains, volcanic rises, andcrustal plateaus make up more than 80% ofthe surface area of Venus. Plains are domi-nant (;65 to 70% by area), reside at lowelevations, and show minor, but broadlydistributed deformation. Nine volcanic rises(2) and five crustal plateaus are similar inplanform shape and size, but differ in topog-raphy, gravity, and surface geology (3, 4).Volcanic rises are dome-like regions (1000-to 2400-km diameters) and contain riftzones, lava flows, and major volcanic edi-

fices (5). Their large gravity anomalies andrelatively low crater densities suggest thatthey are young features (6–8), largely sup-

ported thermally by their buoyant plumesource. Crustal plateaus are steep-sided,flat-topped, quasi-circular regions that hostmost of the planet’s tesserae (complexlydeformed crust). Plateaus have small gravityanomalies consistent with isostatic com-pensation by a thick crust (6, 9–11). As aclass, plateaus are older than rises on thebasis of crater densities (12).

The Origin of Crustal Plateaus

The basis for suggesting that rises and pla-teaus can be explained by mantle plumesand reflect the systematic cooling of Venusis a recent interpretation of the structuralgeology of crustal plateau tesserae indicat-

R. J. Phillips is in the Department of Earth and PlanetarySciences, Washington University, Campus Box 1169,One Brookings Drive, St. Louis, MO 63130, USA. V. L.Hansen is in the Department of Geological Sciences,Southern Methodist University, Dallas, TX 75275, USA.

*To whom correspondence should be addressed.

Fig. 1. Crests of NW-trending folds are cut by NE-trending lens-shaped graben. Ribbons parallel grabenalthough they differ in structural style and geometry (insets). Ribbons (i) are marked by bright and darkpaired lineaments .50 to 100 km long; these lineaments define alternating, flat-topped ridges andnarrow, steep-sided, flat-bottomed troughs. In contrast, graben (ii) host numerous accommodationstructures, they have aspect ratios of 3 to 4 rather than .50, their walls are sloped inward ;60°, andthey are lens-shaped with the widest part of the graben superimposed on fold crests.

SCIENCE z VOL. 279 z 6 MARCH 1998 z www.sciencemag.org1492

ing that the earliest preserved deformationis extensional (13, 14). This contradicts awidely accepted model for the formation ofcrustal plateaus by convective mantledownwelling and secondary crustal thicken-ing (15–17). The downwelling model wassupported by an interpretation of radar im-ages that the oldest preserved deformationalfeatures in these plateaus are contractional,not extensional, in nature (18–20). Geody-namical models of crustal plateau formationby convective downwelling predict that thehigh areas of a plateau should have firstundergone contraction (21, 22), whereas abuoyancy brought about by plumes shouldinitially extend the surface (21, 23).

A key area for understanding the se-quence of deformation on crustal plateaus isthe northeastern part of the crustal plateauOvda Regio. Here, northeast-trending gra-ben do indeed cut the crests of northwest-trending folds, indicating that surface con-traction predated extension (Fig. 1). Closeexamination, however, reveals a third struc-tural feature. “Ribbons” [named for theirlong linear character and high length versuswidth (13)] parallel the late graben yetdiffer in structural style. Ribbon structurescomprise a fabric of steep-sided, alternatingridges and troughs. These structures havebeen recognized as distinct from the late

graben [for example, “steep troughs” (18) or“narrow troughs” (17, 24)], but it was typ-ically assumed that graben-fold temporalrelations applied to ribbons as well. In ad-dition to their difference in structural style,ribbons are much more widely distributedacross individual crustal plateaus, unlike thegraben. Ribbons occur within each of thecrustal plateaus except Phoebe Regio (4,25).

Radar analysis of ribbon structures indi-cates that typical ribbons are steep-sided(.85°), flat-bottomed, shallow (,0.4 km)troughs with V-shaped lateral terminations(Fig. 2) indicating that they result fromopening of tensile fractures in a thin strongbrittle layer above a weaker ductile sub-strate, like a chocolate layer atop a caramelbar (26). Locally ribbons are bounded bysteep extensional shear fractures, represent-ing dominantly brittle deformation of astrong layer above a weak substrate. Thetypical spacing between ribbons of ;1 to 2km and their morphology are consistentwith the inference that ribbons formed fromextension of a thin (,1 km) brittle layer ofthe crust (14). The ribbon floors then markthe depth to a brittle-ductile transition(BDT). These depths are ,0.5 km for ten-sile-fracture ribbons and 0.6 to 1.2 km forshear-fracture ribbons (13). In contrast, the

typical wavelength of crustalplateau folds of ;15 km impliesthat the depth to the BDT when foldsformed was ;6 km (27). Because ribbonsand folds are found together, it is mostreasonable that the ribbons, which reflect ashallow depth of the BDT, formed beforethe folds. If the ribbons formed after thefolds, the folds would not be supported bythe shallow BDT, and they would havedisappeared by viscous relaxation. Howev-er, if ribbons formed first and folds formedlater as the depth to the BDT increasedwith time, the observed geometric charac-teristics of both types of structures could beaccommodated. These mechanical consid-erations are consistent with geometric anal-yses that require ribbons to predate folds(14). The late graben, which formed afterthe folds, likely reflect a still deeper depthto crustal BDT, as evidenced by their spac-ing of .25 km. The complete deforma-tional sequence suggests progressive coolingof an individual crustal plateau with time,and is consistent with plateau formation byan upwelling plume, magmatic injectioninto the crust, solidification, and subse-quent cooling of the rock mass.

The idea that crustal plateaus are genet-ically related to plumes was raised before(28, 29), in which a volcanic rise evolvedinto a crustal plateau by continued subsur-face addition of magmatic material to thecrust overlying the plume. This model wascriticized because crustal plateaus do notshow any remnants of an earlier volcanicrise stage, particularly large shield volcanoes(17–19). The hypothesis also holds thatcrustal plateaus as a class are younger thanvolcanic rises, in conflict with impact craterdensities.

Temporal Constraints

Any plausible geological model of Venusmust satisfy available age constraints,which are based primarily on counts andsizes of impact craters. The average surfaceage of Venus is estimated to be 750 mil-lion years ago (Ma) (30). Tesserae, whichmake up most of the crustal plateaus, havean average age of 900 6 220 Ma (31).Volcanic rises have large gravity anoma-lies, implying that they are dynamicallysupported at depth today and thus areyoung. Additionally, features associatedwith volcanic rises (for example, large vol-canoes) are younger than the global aver-age surface age by up to a factor of 4 (8).In some cases, however, the average sur-face age of a rise is not young becausemuch of the surface is composed of uplift-ed old plains (for example, Beta Regio).

The volcanic plains of Venus are rough-ly divisible into four geomorphic units (32,

12°E59°N

25 km

5 km

Fig. 2. North-trending rib-bons (paired dark and lightlineaments) track acrossE-trending fold crests inFortuna Tessera. The ap-parent deflection of the rib-bons across the crest ofthe folds toward the radar(left) is the result of radarforeshortening, an imagingartifact; parallelism of op-posite ribbon troughslopes across the folds in-dicates that the slopes arevery steep, .85° (14). In-sets show an enlargementof the region in the box;ribbon troughs are markedby tic marks, ridges arebetween the troughs. NoteV-shaped joining of wallson opposite sides of thetroughs.

www.sciencemag.org z SCIENCE z VOL. 279 z 6 MARCH 1998 1493

33) that overlap in time and span a totalage range of more than 1000 to 600 Ma(30). The two youngest plains units arespatially associated with chasmata (rifts)and centers of young constructional volca-nism (32). More generally, the two young-est plains units can be combined into asingle unit of “young plains” (33% of totalplains area) with an average age of 675 650 Ma, and the two other units lumped into“old plains” (34) with an average age of975 6 50 Ma. This division is visuallyobvious on a global map of crater density.

Global Geological Model

The geological and temporal relations sum-marized above can be explained if crustalplateaus and volcanic rises both formedabove plumes, but at distinctly differenttimes, with different lithospheric thickness-es, and hence within different magmatic en-vironments because of the depth of melting(Fig. 3). We suggest that plateaus formedearly when the lithosphere was thin (mobileor thin-lid regime) during a time of prolificmagmatism; rises formed later in thick litho-sphere (stagnant- or thick-lid regime) whenmagmatic activity was relatively low. Thetransition in magmatic rates could have beenbrought about by a change in convectivestyle in the mantle attributable simply tosecular cooling of the planet. We suggestthat in addition to centralized crustal plateau

magmatism, the thin-lid regime also per-mitted regional eruptions of volcanic flowsto form most of the existing plains. Asfocused plumes rising from the core-mantleboundary were undergoing extensive partialmelting in the upper mantle to form crustalplateaus, internally heated convection inthe mantle caused diffuse upward motion oflarge volumes of mantle (35), triggeringspatially widespread partial melting andplains volcanism as a result of pressure-re-lease melting under a thin lithosphere.

During Venus’ thin-lid period, a processakin to plate tectonics may have developed,although delamination of the uppermostmantle beneath the crust could insteadhave served to recycle cold material to thedeeper mantle (Fig. 3). During plume up-welling in this period, the crust heats rap-idly by injection of magma; it uplifts, thick-ens, then cools and subsides. Ribbon fabricswould be expected to form first during up-lift, when the BDT was shallow and tensilestresses were high. Gravitational sliding andlimited contraction along the plateau mar-gins would have thickened the crust inthese regions. As the hot, thickened, crusteventually cooled; local slumping of thepreviously inflated crust led to local, limit-ed, disorganized shortening at the surface,and the BDT deepened. Eventually, a pla-teau is left with thick isostatically compen-sated crust, which is unstable and so, withtime, collapses into the surrounding low-

lands. This model implies that the existingcrustal plateaus formed relatively late in thethin lithospheric lid phase and were cap-tured and preserved within the thickeninglithosphere. We propose that earlier formedcrustal plateaus are locally preserved aslarge tessera inliers (3, 4), many of whichhost ribbon terrain structures (17, 36). Af-ter the transition to a thick lid, largeamounts of melting in the mantle is nolonger possible, so volcanic rises, ratherthan crustal plateaus, form on the stronger,thicker lithosphere. Lacking, a crustal root,volcanic rises must be supported dynamical-ly by plumes and will collapse as the plumeebbs. The stress state of the lithosphere alsochanges, so that rifts, and coronae (37)associated with them, become the prevalenttectonic style. Plains volcanism continues,but at a greatly reduced rate (33).

Our model and age estimates of the Ve-nus crust provide a prediction of how rap-idly the lithosphere thickened. We placethe switch from mobile (thin) to stagnant(thick) lid at 700 Ma under the suppositionthat some of the young plains formed duringthe transition period from thin to thick lidwhen the stress state in the lithospherefundamentally changed and magmatism wasdeclining. The absence of extensive plainsyounger than 600 Ma might suggest thatformation of plains ceased quickly, althoughwe do not know the true age range of plains.The limited number of embayed craters onthe plains may imply that volcanism tailedoff rather quickly, perhaps over 100 My(12), but this evidence is weak. Anotherargument is that if the lithosphere thick-ened gradually, then, according to our hy-pothesis, we should observe a number ofstructures that are transitional betweencrustal plateaus and volcanic rises, yet wefind only one candidate feature: PhoebeRegio is characterized by a complex patternof extensional graben and lack of ribbons(4, 13).

A change in mantle convective style at;700 Ma could have led to a rapid increasein the thickness of the lithosphere. Becausethe viscosity contrast between the surfaceand the interior of a terrestrial planet islarge, a thick, stagnant lithospheric lidshould form (38). An exception to thisscenario occurs if, like the Earth, a planet isable to recycle a portion of its strong litho-sphere because it can be broken (that is, athin- or mobile-lid regime; Earth’s versionof this is plate tectonics). As a planet coolsand radiogenic heat production wanes, con-vection in the mantle slows and the in-duced stress on the overlying lithospherediminishes. Heat flux into the lithospherealso decreases, so the lithosphere cools andstrengthens and becomes harder to break.Eventually, recycling comes to an end, and

collapsed crustal

plateaus

plains volcanism

crust

mantle portion of thermal lithosphere

convecting mantle

ancient crustalplateaus

core-mantlethermal

boundary layer

core

plains basins

volcanic rises

crustal plateaus

coronae andchasmata? to 0.7Ga

~ 0.5 Ga to present

lithospheric recycling

Ga1.6 1.2 0.8 0.4 0.0

120

100

80

60

40

20

km

Fig. 3. Cartoon illustrating the major tectonic and magmatic processes operating in the thin (mobile)-lidregime (.0.7 Ga) and in the thick (stagnant)-lid regime (;0.5 Ga to present). The boxed inset showsthermal lithosphere thickness (km) versus age as calculated from a thermal model. In the thin-lid era,lithospheric recycling was associated with mantle downwellings whose surface expression includedridge belts. In the thick-lid era, lithospheric extension over a heating mantle produces corona chains andchasmata, which mark broad diffusive mantle upwellings; the topographic basins of the plains reflectmantle downwellings (3).

SCIENCE z VOL. 279 z 6 MARCH 1998 z www.sciencemag.org1494

the lithosphere rapidly thickens and be-comes stagnant. This switch from a mobileto stagnant lithospheric lid is a natural con-sequence of terrestrial planetary evolution.The question is whether a stagnant lid isstable or whether it can collapse into avigorous but transient episode of litho-spheric recycling (subduction) followed byreformation of a stagnant lid (39).

Some numerical models of mantle con-vection show repeated episodes of stagnantlid collapse, each followed rapidly by ashort-lived episode of lithospheric subduc-tion (39). Similarly, one model of the evo-lution of Venus (16, 40), based on inferenc-es from the cratering record (41), is that atransient subduction event led to wide-spread volcanism and rapid global resurfac-ing of the planet. However, the crateringrecord provides information only into theperiod of resurfacing decline and cannot beused to distinguish between a punctuatedresurfacing event and the end of a sustainedperiod of resurfacing. We suggested abovethat the major plains units span about 400My and extend back in time from the tran-sition from a thin to a thick lid. We cannotsay whether there were transient eventsbefore this period, or whether in the futurethe present thick lid will collapse. Addi-tionally, some models (16, 40) predict thatthe present lithospheric thickness is about300 km; this value is sufficiently thick toprohibit magmatic activity, contrary to ob-servation (33).

Herrick (42) proposed that a mobile-lidphase ended on Venus in the last billionyears. His model differs from ours in twoways. First, he suggested that tesserae recordthe mobile-lid deformation and are globallypervasive, whereas we argue that mosttesserae are associated with crustal plateausor remnants of them, as supported by geo-logic mapping (25). Secondly, Herrick ar-gued that the volcanic plains were rapidlyemplaced at the end of the mobile-lid eraafter all tesserae had formed, while we sug-gest that plains deposition was an ongoingprocess during the thin-lid era. The westernmargin of Tellus Regio preserves ;1000 kmof a deformation front in which crustalplateau tesserae grade into preexistingplains units, so tesserae are not everywhereolder than plains, as has been suggested (12,17, 43). Elsewhere, many of the youngestvolcanic plains units onlap tesserae at crust-al plateau boundaries. Thus in most casesthe youngest portion of volcanic plainsformed locally after adjacent crustal pla-teaus, although we argue that both typesformed broadly at the same time.

During the formation of the voluminousplains, outgassing of volatiles should haveenhanced the atmospheric greenhouse ofVenus and thereby increased the surface

temperature. In turn, higher surface temper-atures would have decreased the heat fluxfrom the lithosphere and increased interiortemperatures and the amount of melting.Enhanced magmatism would then providemore volatiles to the atmosphere. Such afeedback mechanism would imply that theinterior and climatic processes of Venuswere tightly coupled during the thin-lidphase. Surface temperatures could have ap-proached 1000 K (44) during thin-lid time.

Computational Models

To further test the geodynamical plausibil-ity of our interpretations, we consideredresults from a numerical simulation of thethermal evolution of Venus and of the gen-eration of magmas in the mantle. We com-puted a thermal evolution using a one-dimensional parameterized convection ap-proach (45, 46). The thermal structure ofthe mantle and of plumes rising off thecore-mantle boundary, along with mini-mum depth of melting controlled by litho-sphere thickness, determine melting behav-ior. Temperatures in the convecting mantleand in the core are determined by using theconservation of thermal energy along with arelationship between the Nusselt numberand the thermal Rayleigh number (47). Theparameters in the Nusselt number–Rayleighnumber relationship are changed com-mencing at 700 Ma; the Nusselt numberdecreases and the lithosphere thickens inresponse. In thin-lid convection, heattransfer is dominated by cooling of themantle with recycled lithosphere; the char-acteristic conversion time to thick-lid be-havior of the Nusselt number–Rayleighnumber relationship is equated to the ther-mal assimilation time of the subducted pre-change lithosphere. The time scale of litho-spheric thickening is controlled by the ther-mal equilibrium time of the new lithospheremodeled as a conductive slab. The results of

a specific thermal model calcu-lation yield time-dependent esti-mates of mantle and core temperatures,lithospheric thickness, and heat flux, pro-viding the basic information for partial meltcalculations. We included the effects of anenhanced greenhouse by using a 1000 Ksurface temperature during thin-lid timeand then decreasing it to 740 K during the50 My after a thick lid started to form.

Plume temperature was obtained from thecore-mantle thermal boundary layer, weight-ed 75% and 25% between the solution tem-peratures obtained at the bottom and top ofthe boundary layer, respectively (48). Toobtain a temperature for melting calcula-tions near the top of the mantle, the plumewas cooled adiabatically and also assumed tolose diffusively a fraction, fp, of its initialtemperature. The temperature for melting ofdiffuse mantle upwellings to produce plainsvolcanism is the upper mantle temperaturepossibly increased by a small amount DTp.The two temperatures (plumes, plains) ob-tained are converted to potential tempera-tures by adiabatically reducing them to theirzero pressure values because potential tem-perature is used in determining melt fraction.We assume that the lithosphere (thickness,H) was thermally eroded and stretched sothat the effective depth for melting isHDLplu for plumes and HDLupw for upwell-ings. These values establish minimum melt-ing depths. Given potential temperature andlithostatic pressure at the base of the erodedand stretched lithosphere, partial melt frac-tion, or equivalent melt column height, canbe calculated for both plumes and plains(49). The numerical solution used as inputfor a partial melting calculation is called a“computational thermal model” (CTM), andthe result of a partial melting calculation isreferred to as a “computational melting mod-el” (CMM).

There are a relatively large number offree parameters in the CTM, but Table 1

Table 1. Constraints for thermal and magmatic modeling. CP, crustal plateau; VR, volcanic rise; Ga,billions of years ago.

Observable Interpretation or estimation Implication

Ribbons CP earliest strain eventextensional

Plume origin for CPs

Gravity CP have thick crustal rootsVR dynamically supported

Magmatism played a strongerrole in CP formation

Heat flow 36 to 55 mW m22 Present lithosphere thicknessis Earth-like

Phoebe Regio Only candidate for transitionfrom CP to VR

Lithosphere thickened quickly

Global geological modelPlumes First created CP then VR Melting in plumes declined at

0.6 to 0.7 GaPlains Plains generation extensive

before 0.7 Ga; meagerduring present thick-lid era

Melting in mantle declined at0.6 to 0.7 Ga

www.sciencemag.org z SCIENCE z VOL. 279 z 6 MARCH 1998 1495

shows that the combination of geologicaland geophysical constraints, plus our globalgeological model, provide a number ofguidelines for selecting parameter values(50). It was not our intent to use the con-straints in Table 1 to estimate parameters ofthe CTM; rather, we sought geologicallyand geophysically plausible sets of parame-ters to satisfy the constraints listed. Thusthe CTM was tuned to the geological andgeophysical constraints described above.Subsequently we investigated other predic-tive properties of the CTM for the param-eter values adopted, including the crustalplateau formation rate, the plains emplace-ment rate, and the behavior with time ofthe depth to the BDT in crustal plateaus.Because plains, crustal plateaus, and volca-nic rises areally represent most of Venus,our results have global implications.

The nominal computational model (51)satisfies the constraints of Table 1, includ-ing a present-day surface heat flow of 44mW m22. At the convective switch, thelithosphere thickens from ;40 km to over100 km over ;200 My (Fig. 3). At the sametime the upper mantle temperature increas-es while the rate of cooling of the outer coredecreases (Fig. 4A). This evolution impliesthat the partial melting histories of plainsand plumes will be quite different. A suiteof models varying the lithosphere erosion-stretching parameter, HDLupw, shows that abasic feature of plains generation is, afterthe switch, a decrease in melt generationdue to the rapidly thickening lithospherefollowed by a partial recovery due to theincreasing temperature of the upper mantle(Fig. 4B). Depending on the choice ofHDLupw, partial melting can be completelyeliminated or decreased by only a moderateamount (52). The nominal model has amodest recovery in plains volcanism, con-sistent with the imposed constraints of Ta-ble 1 (Fig. 4C). Also calculated is a plumecurve (Fig. 4C), where melt generation de-creases to a relatively constant value; herethe increase in lithospheric thickness dom-inates melt behavior because the core cool-ing rate, which controls plume temperature,changes only modestly. The decrease inmelt generation is nearly a factor of 20,consistent with the generation of volcanicrises, not crustal plateaus, after the switch.

Our calculations (Fig. 4) indicate thepotential for pressure-release partial meltingwhen hot mantle material is brought nearthe base of the lithosphere, but they do notprovide rates of melt generation. The plumerate can be estimated from the amount ofmodel-calculated heat supplied by the coreto the core-mantle thermal boundary layer,where plumes originate (53). In the billionyears preceding the switch, the CMM pre-dicts (CTM parameters were not tuned forthis particular result) that enough partialmelt by volume is generated to produceabout 13 features the size of the crustalplateau Tellus Regio (54). This amountwould account for the five observed crustalplateaus in about the same interval of geo-logical history. Given this rate, and thepossibility of at most one observed geologicstructure, Phoebe Regio, that could be in-termediate between crustal plateaus andvolcanic rises, the expected thickening timeof the lithosphere is a few hundred millionyears. This result supports our hypothesisthat the lithosphere must have thickenedrapidly compared to that expected from theslow secular cooling brought about by a lossof heat-producing parent isotopes (withouta corresponding change in convectivestyle). Because the CTM shows that achange in mantle convective style has only

a small effect on core cooling, it also followsthat plume-derived crustal plateaus couldnot have all formed during a single punctu-ated event of subduction (16).

Broad upwellings associated with mantleconvection should produce widespread, ac-tive plains volcanism in the thin-lid era.According to the CMM, this process pro-duces a volume partial melt fraction, w, of21% at 1.7 Ga decreasing to 17% at 0.7 Ga.The volume per unit area of melt producedfor plains emplacement is the product of thediffuse upwelling velocity, vu, and w. If vu iseven as small as 0.1% of the convectionboundary layer velocities obtained from theCTM, then vertical plains emplacementrates vary from 10 km per billion years at1.7 Ga to 5 km per billion years at 0.7 Ga.Higher velocities may imply significantcrustal recycling, but the major point is thata plausible set of model parameters predictsthat volcanic plains emplacement was adominant, ongoing process during thin-lidtime and not a punctuated event. Given therarity of crustal plateau genesis, then, asobserved, exposed volcanic plains should bealmost everywhere younger than crustalplateaus. This does not, however, implythat plains emplacement was a single eventthat followed all crustal plateau formation,contrary to other viewpoints (12, 17, 43).

Ribbon Formation

The development of tensile ribbons placessevere constraints on the thermal environ-ment of crustal plateaus. The shallow crusthad to be free of preexisting faults andthroughgoing fractures, which would havebeen reactivated at stresses lower than thetensile strength of the crust (55, 56). Vastareas in crustal plateaus must have beenmechanically suited to ribbon formation.Either a plateau crust was not previouslydeformed, or it had been annealed by solid-state creep or magma injection. Further-more, the crust must have been hot enough,long enough, to sustain tensile failure atshallow depths. The magmatic origin ofcrustal plateaus may provide the proper en-vironment; injection of magma into frac-tures could have created a mechanicallyhomogeneous environment in which toform ribbons. At a surface temperature of740 K, model calculations (57) show thatthe BDT would rapidly deepen as the crustcooled; shear failure would be expected at adepth of 1 km after only about 20,000 years(58). At a surface temperature of 1000 K,this time increases by a factor of 15, and theshear fracture ribbons could be expected tohave formed within 1 My. This time ap-proaches the duration of large terrestrialflood basalt events, which have been linkedto substantial partial melting in mantle

B

A

-75

-50

-25

0

25

50

Tem

pera

ture

(K

)

Age (Ga)

Mel

t Hei

ght (

km)

10-2

10-1

100

101

102

Mel

t Hei

ght (

km)

1.6 1.2 0.8 0.4 0.010

-1

100

101

102

C

-100

plumes

plains

mantle

core

Fig. 4. (A) CTM-computed temperatures as afunction of age in the uppermost convecting man-tle (blue) and at the top of the iron core (red) ref-erenced to their values at 1.6 Ga. (B) A suite ofCMM models for the height of plains-forming meltcolumn as a function of lithospheric penetration,DLupw. Curves from top to bottom are a range inDLupw from 0.70 to 1.05 in increments of 0.05. (C)Melting curves for nominal model (51) of plains(blue) and plumes (red).

SCIENCE z VOL. 279 z 6 MARCH 1998 z www.sciencemag.org1496

plume heads (59). The most important mes-sage from the ribbons may be that the cli-mate and internal history of Venus werestrongly coupled throughout much of itshistory.

REFERENCES AND NOTES___________________________

1. Lithosphere is defined as the strong outer shell of aterrestrial planet that transports heat to the surfaceby conduction.

2. S. E. Smrekar, W. S. Kiefer, E. R. Stofan, in Venus II,S. W. Bougher, D. M. Hunten, R. J. Phillips, Eds.(Univ. of Arizona Press, Tucson, 1997 ), pp. 845–878.

3. R. J. Phillips and V. L. Hansen, Annu. Rev. EarthPlanet. Sci. 22, 597 (1994).

4. V. L. Hansen, J. J. Willis, W. B. Banerdt, in (2), pp.797–844.

5. E. R. Stofan et al., Geol. Soc. Am. Bull. 101, 143(1989); D. A. Senske, G. G. Schaber, E. R. Stofan, J.Geophys. Res. 97, 13395 (1992).

6. S. E. Smrekar and R. J. Phillips, Earth Planet. Sci.Lett. 107, 582 (1991).

7. R. E. Grimm and R. J. Phillips, J. Geophys. Res. 97,16035 (1992); S. E. Smrekar, Icarus 112, 2 (1994).

8. N. Namiki and S. C. Solomon, Science 265, 929(1994); M. Price and J. Suppe, Nature 372, 756(1994); R. J. Phillips and N. R. Izenberg, Geophys.Res. Lett. 22, 1517 (1995); M. H. Price, G. Watson,J. Suppe, C. Brankman, J. Geophys. Res. 101, 4657(1996).

9. R. R. Herrick and R. J. Phillips, J. Geophys. Res. 97,16017 (1992).

10. R. E. Grimm, Icarus 112, 89 (1994); M. Simons, S. C.Solomon, B. H. Hager, Geophys. J. Int. 131, 24(1997 ).

11. R. E. Grimm and P. C. Hess, in (2), pp.1205–1244.12. A. T. Basilevsky, J. W. Head, G. G. Schaber, R. G.

Strom, in (2), pp.1047–1084.13. V. L. Hansen and J. J. Willis, Icarus 123, 296 (1996).14. iiii, ibid., in press.15. S. C. Solomon, Phys. Today 46, 48 (1993).16. D. L. Turcotte, J. Geophys. Res. 98, 17061 (1993).17. M. A. Ivanov and J. W. Head, ibid. 101, 14861

(1996).18. D. L. Bindschadler G. Schubert, W. M. Kaula, ibid.

97, 13495 (1992).19. D. L. Bindschadler A. DeCharon, K. K. Beratan, S. E.

Smrekar, J. W. Head, ibid., p. 13563.20. D. L. Bindschadler, Rev. Geophys. (Suppl.) 33, 459

(1995).21. D. L. Bindschadler and E. M. Parmentier, J. Geo-

phys. Res. 95, 2132 (1990).22. A. Lenardic, W. M. Kaula, D. L. Bindschadler, Geo-

phys. Res. Lett. 18, 2209 (1991).23. R. J Phillips, J. Geophys. Res. 95 1301 (1990).24. J. W. Head, Lunar Planet Sci. XXVI, 579 (1995).25. R. R. Ghent and V. L. Hansen, ibid. XXVIII, 409

(1997 ).26. The type ribbons are located in southeastern Fortuna

Tessera (14).27. C. D. Brown and R. E. Grimm, Earth Planet. Sci. Lett.

147, 1 (1997 ).28. R. R. Herrick and R. J. Phillips, Geophys. Res. Lett.

17, 2129 (1990).29. R. J. Phillips, R. E. Grimm, M. C. Malin, Science 252,

651 (1991).30. W. B. McKinnon, K. J. Zahnle, B. I. Ivanov, H. J.

Melosh, in (2) pp. 969–1014. Age estimates of aspecific geologic unit are obtained by scaling its cra-ter density to the global mean crater density.

31. Only craters with diameters $32 km are counted toavoid a potential observational bias [M. A. Ivanov andA. T. Basilevsky, Geophys. Res. Lett. 20, 2579(1993)]. This estimate averages, across all of thetesserae, the oldest surviving crater on each tessera(the “retention age”). This age is less than the aver-age formation age of the tesserae themselves. Theerrors used here and elsewhere are 1s values basedon counting statistics and do not take into accountthe uncertainty in the 750 Ma mean age value (30).

32. M. Price, Proc. Lunar Planet. Sci. Conf. 26, 1143(1995); in preparation.

33. K. L. Tanaka, D. A. Senske, M. Price, R. L. Kirk, in (2),pp. 667–694.

34. S. A. Hauck II, R. J. Phillips, M. H. Price, J. Geo-phys. Res., in press.

35. G. T. Jarvis and W. R. Peltier, in Mantle Convection:Plate Tectonics and Global Dynamics, W. R. Peltier,Ed. (Gordon & Breach, New York, 1989) pp. 479–593.

36. J. W. Head, Lunar Planet. Sci. XXVI, 573 (1995).Models that progressively flood crustal plateaus re-veal characteristics typical of large tessera inliers,including arcuate marginal traces.

37. Coronae on Venus are “ . . . any circular to elongatestructure defined primarily by an annulus of concen-tric fractures and ridges.” They occur most often inassociation with chasmata (rifts) [E. R. Stofan, V. E.Hamilton, D. M. Janes, S. E. Smrekar, in (2), pp.931–965].

38. V. S. Solomatov, Phys. Fluids 7, 266 (1995); G.Schubert, V. S. Solomatov, P. J. Tackley, D. L. Tur-cotte, in (2), pp.1245–1288; L.-N. Moresi and V. S.Solomatov, Phys. Fluids 7, 2154 (1995); V. S. Solo-matov and L.-N. Moresi, J. Geophys. Res. 101,4737 (1996).

39. L. Moresi and V. Solomatov, Geophys. J. Inter., inpress.

40. D. L. Turcotte, J. Geophys. Res. 100, 16931 (1995).41. G. G. Schaber et al., ibid. 97, 13257 (1992); R. G.

Strom, G. G. Schaber, D. D. Dawson, ibid. 99,10899 (1994).

42. R. R. Herrick, Geology 22, 703 (1994).43. A. T. Basilevsky and J. W. Head, Earth Moon Planets

66, 285 (1995).44. Model calculations show that only modest increases

(15%) of H2O and SO2 in the Venus atmosphere willdrive the surface temperature to 1000 K [M. A. Bul-lock and D. H. Grinspoon, J. Geophys. Res. 101,7521 (1996)].

45. R. J. Phillips and M. C. Malin, in Venus, D. M. Hunten,L. Colin, T. M. Donahue, Eds. (Univ. of Arizona Press,Tucson, 1983), pp. 159–214.

46. D. J. Stevenson, T. Spohn, G. Schubert, Icarus 54,466 (1983).

47. The Nusselt number, Nu, is the ratio of conductiveplus convective heat transport to conductive heattransport alone in a fluid system. The Rayleigh num-ber, Ra, is the ratio of buoyancy forces to viscousforces in a fluid. These quantities are related througha power exponent b and a viscosity length scale pwith power dependence a: Nu ; p aRab. The valueof b is different for thin-lid and thick-lid convection(and a [ 0 for thin-lid convection). Two coupledordinary differential equations are integrated in timeto obtain upper mantle (at base of lithospheric ther-mal boundary layer) and outermost core tempera-tures. Core energy balance includes the effects ofsolid inner core formation. Heat sources in the man-tle are fractionated into a crust with time, and theheat flux out of the convecting mantle depends lin-early on Nu.

48. The boundary layer bottom temperature is the out-ermost core temperature, and the top temperature is

obtained by adiabatic extrapolation ofthe upper mantle solution temperature.

49. D. McKenzie, J. Petrol. 25, 713 (1984); D. McKenzieand M. J. Bickle, ibid. 29, 625 (1988).

50. The heat flow estimate in Table 1 is obtained fromflexural modeling of a volcanic rise with a correctionto the background value [R. J. Phillips et al., in (2), pp.1163–1204].

51. Parameter values are f p 5 0.05, DLplu 5 0.80,DLupw 5 0.90, DTp 5 0 K, activation energy for creepdivided by gas constant 5 50,000 K, thermal diffu-sivity 5 5 x 1027 m2 s21, b (thin lid) 5 0.33, and b(thick lid) 5 0.30.

52. If these calculations are carried forward in time,partial melting reaches a peak and then starts toslowly decrease due to the loss of heat-producingradiogenic parent isotopes in the mantle. This isalso the reason for the slope prior to the switch.

53. N. H. Sleep [J. Geophys. Res. 95, 6715 (1990)] pre-sented a simple relationship between buoyancy fluxin a plume (B in kg s21) and heat flow (H in W): whereCp is specific heat and a is volume coefficient ofthermal expansion. Using this, the partial melt mass,Mpm, over a time interval, t, can be related to heatflow from the core, qc, by Mpm 5 aX4pRc

2qct/Cp,where X is the mass fraction of partial melt and Rc iscore radius. This can be converted to volume of rock,assuming values for density and contraction. Thesubsequent mass is then divided by the root mass ofa Tellus-Regio–sized crustal plateau (2 x 106 km2, 1km average height above its surroundings) using anisostatic assumption. This yields the number of Tel-lus-sized crustal plateaus generated by plumes overthe interval t.

54. This is obtained from an average melt volume frac-tion of 20% and an average core heat flux of 22 mWm22 during the interval 1.7 Ga to 0.7 Ga.

55. R. H. Sibson, J. Struct. Geol. 7, 751 (1985); M. P.Golombek and W. B. Banerdt, Icarus 68, 252(1986).

56. M. P. Golombek and W. B. Banerdt, ibid. 83, 441(1990).

57. This result is obtained by constructing a yieldstrength envelope [W. F. Brace and D. L. Kohlstedt,J. Geophys. Res. 85, 6248 (1980)] using a dry dia-base flow law [S. J. Mackwell, M. E. Zimmerman,D. L. Kohlstedt, D. S. Scherber, in Proc. of the 35thU.S. Symposium on Rock Mechanics, J. J. K. Dae-men and R. A. Shultz, Eds. (A. A. Balkema, Rotter-dam, 1995), pp. 207–214] and 10% strain in 0.1 My.The temperature profile used was for a cooling half-space model with a surface temperature of 740 or1000 K and an initial temperature determined bymixing melt temperatures from the CMM with crustaltemperatures from the CTM calculated with the ap-propriate surface temperature.

58. J. C. Jaeger and N. G. W. Cook, Fundamentals ofRock Mechanics (Chapman & Hall, New York, 1979).The transition depth for tensile to shear failure isdependent on the choice of tensile yield stress. Weused a value of 10 MPa, which is within the rangeestimated for intact basalts [R. A. Schultz, J. Geo-phys. Res. 95, 10883 (1993)].

59. C. G. Farnetani and M. A. Richards, J. Geophys.Res. 99, 13813 (1994).

60. This work was supported by NASA’s Planetary Ge-ology and Geophysics programs under GrantsNAGW-3024 to Washington University and NAGW-2915 to Southern Methodist University. D. Brownand S. Hauck provided helpful reviews of earlierdrafts of this manuscript. We thank M. Zuber andtwo additional reviewers for comments. J. Goodgeprovided a review of the revised manuscript.

www.sciencemag.org z SCIENCE z VOL. 279 z 6 MARCH 1998 1497