October 2007 LIP of the Month

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Large Igneous Provinces (LIPs) on Venus

Vicki L. Hansen

Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812

Email: vhansen@d.umn.edu; Tel: (218)-726-6211; Fax (218)-726-8275


Venus, Earth’s sister planet lacks plate tectonics, yet hosts numerous examples of large igneous provinces (LIPs); thus, Venus may provide critical clues to terrestrial LIP formation processes. Volcanic rises, which represent contemporary plume signatures on thick lithosphere, include three morphologically distinct groups, rift-, volcano- and coronae-dominated rises; each group may represent differences in lithospheric properties and/or relative stages of formation. Four large coronae, distinguishable from other coronae based on size, are morphologically quite different from one another, and may record vastly different processes of formation. Artemis likely represents the surface manifestation of a plume on thin lithosphere; Heng-O’s evolution is unclear and could represent either endogenic or exogenic driven processes. Quetzalpetlatal requires more mapping; preliminary work favors Quetzalpetlatal’s formation above a plume, however Quetzalpetlatal could also reflect the culmination of several distinct events. Atahensik, the only large coronae to lie within a coronae chain represents a large endogenic diapiric structure, though the nature of the diapiric buoyancy is unconstrained. Crustal plateaus, previously variably interpreted as endogenic features forming above deep mantle downwellings or deep mantle plumes, likely represent exogenic features. Their distinctive tectonic fabric with areally extensive but structurally localized lava flooding requires a huge amount of lava to have existed at one time. Crustal plateaus may represent the surface scum of huge lava ponds formed by massive partial melting in the mantle due to the impact of a large bolide with ancient thin lithosphere. 


Vigorous debate regarding the origin of large igneous provinces, LIPs, centers on whether LIPs result from plate-related processes (e.g., Anderson and Nathlan, 2005), or deep-mantle plumes that likely track to the core-mantle boundary (e.g., Campbell and Griffiths, 1990; 1992; Eldholm and Coffin, 2000; Ernst et al., 2005). Imagine a planet similar to Earth that displays large tectonomagmatic provinces, but lacks plate tectonic processes.  Venus fits the bill. Venus is expected to have a heat budget similar to Earth given similarities in size, age, density and presumed composition; thus Venus should require mechanisms through which heat is transferred from core (or mantle) to crust.  However, Venus lacks plate tectonic processes (Solomon et al., 1992; Phillips and Hansen, 1994; Simons et al., 1994, 1997); therefore plate-related hypotheses for LIPs are not viable on Venus. In addition, a lack of plate tectonic processes, and associated recycling of large tracts of a planet’s surface, might result in the preservation of extensive surfaces that preserve rich geologic histories, which could in turn record critical clues about LIP formation. Venus also affords a view of primary structural morphology of dynamical processes (endogenic and exogenic) unequaled by any planet. Venus lacks a hydrologic cycle (being much too hot for surface water to survive); thus Venus’ surface is free of the effects of erosional and depositional processes. There are, however, a few challenges: travel to Venus is costly and the field environment extreme. Fortunately, widely available NASA Magellan Mission data provide incredible global and high-resolution views that allow for first-order geologic analysis, coupled with subsurface constraints based on gravity-topography data.  Venus ‘fieldwork’ only requires computer access, making data collection, data analysis, and hypothesis testing widely accessible.


Venus and Earth share similarities, yet they also have profound differences. Venus, 0.72 AU from the Sun, is 95 and 81.5 percent Earth’s size and mass, respectively. Venus II (Bouger et al., 1997) provides an excellent comprehensive introduction to Venus. Venus and Earth share similar bulk composition and heat producing elements; Venus’ surface is likely basaltic. Slow retrograde motion makes a Venus day longer than its year, a factor that may contribute to Venus’ lack of a magnetic field. Venus’ surface pressure (~95 bars) and temperature (~475°C) vary significantly from Earth, although its atmosphere composition (96% CO2, 3.5% N2 and 0.5% H20, H2SO4, HCl and HF) might be similar to that of early Earth (Lecuyer et al., 2000). Venus’ thick atmosphere results in negligible diurnal temperature variations, and an enhanced greenhouse prohibiting a terrestrial water cycle. Venus lacks obvious evidence of: extensive sedimentary layers, weathering, erosion, sediment transport and deposition processes that play dominate roles in shaping Earth’s surface. Although Venus is presently ultra-dry, the past role of water is unknown. Isotopic data are consistent with, but do not require, extensive reservoirs of water >1 billion years ago (Donahue and Russell, 1997; Donahue et al., 1997; Donahue, 1999; Lecuyer et al., 2000; Hunten, 2002). The present lack of water renders contemporary crustal rock orders of magnitude stronger than terrestrial counterparts, even given Venus’ elevated surface temperature (Mackwell et al., 1998). 

Most workers assume that Venus’ mantle is similar in composition and temperature to Earth’s mantle, with similar viscosity and a strong temperature-dependent viscosity profile. However, some workers consider Venus’ mantle to be stiffer than Earth’s due to presumed drier conditions (e.g., 1020 Pa s, Nimmo and McKenzie, 1998). Volatiles are of course important in understanding Venus’ interior—particularly with regard to viscosity; however, interior volatile values and compositions are currently unconstrained. Lack (presence) of volatiles will increase (decrease) strength and increase (decrease) the mantle solidus. Large viscosity contrasts are likely across thermal boundary layers: notably across the lithosphere and core-mantle boundaries. In contrast to Earth however Venus lacks a low viscosity asthenosphere, likely a contributing factor to Venus’ lack of plate tectonic processes. In discussion of Venus one must be ever mindful of Venus-Earth differences, as well as assumptions and unknowns for both planets. Venus provides a superb sister with which to construct comparative experiments aimed at understanding terrestrial planet evolution.

The NASA Magellan mission returned four remote global data sets: altimetry, synthetic aperture radar (SAR) imagery, emissivity, and gravity data (Ford and Pettengill, 1992; Ford et al., 1993). Emissivity is chiefly controlled by dielectric permittivity and surface roughness (Pettengill et al., 1992). Gravity data can resolve features >400 km, and provides clues to subsurface architecture, although interpretations are notably non-unique. Altimetry has a spatial resolution of 8 km (along-track) by 20 km (across-track), and vertical resolution ~30 m. SAR data (~120m/pixel available via the web at http://pdsmaps.wr.usgs.gov/maps.html), provides the highest resolution view of Venus, and allows for geomorphic and geological interpretations, including geologic surface histories. SAR and altimetry data can be digitally combined to construct synthetic stereo, 3D, views of the surface (Kirk et al., 1992). Cautions for interpretation of geologic features and histories are outlined in a variety of references (Wilhelms, 1990; Ford et al., 1993; Tanaka et al., 1994; Zimbelman, 2001; Hansen, 2000).

Venus is divisible into three topographic provinces: lowlands, highlands and mesolands. The lowlands (70-80% of the surface) include long-wavelength (thousands of km) low amplitude basins (<1 km), marked by relatively smooth low strain regions with local belts of concentrated deformation (Banerdt et al., 1997). The highlands (8-10% of the surface) host volcanic rises, crustal plateaus, and the unique feature Ishtar Terra (Hansen et al., 1997). Mesolands lie between the lowlands and highlands, and host many of Venus’ coronae, (60-2600 km diameter quasi-circular tectonomagmatic features) and spatially associated chasmata (troughs). About 970 impact craters (1-270 km diameter) pepper the surface with a spatial distribution indistinguishable from random (Phillips et al., 1992; Schaber et al., 1992, Herrick et al., 1997; Hauck et al., 1998).  A variety of volcanic landforms ranging in size from km to 100s of km occur across the surface. Lava flows, with areal extent 100s to 1000s of km, are commonly associated with volcanoes, coronae, and fractures (Head et al., 1992; Crumpler et al. 1997). Volcanic shields (1-20 km diameter), occur in shield fields (<300 km diameter regions) and as ‘shield terrain’ distributed across millions of km2 (Guest et al., 1992; Crumpler et al., 1997; Hansen, 2005).  Fluid cut channels, or canali, extending 10s or 1000s of km trace across the surface (Baker et al., 1997).  Volcanic forms are consistent with basaltic compositions (e.g., Bridges, 1995, 1997; Sakimoto and Zuber, 1995; Stofan et al., 2000).

The recently proposed revised definition of LIPs (Bryan and Ernst, 2007) concerns composition, size and rapid emplacement: Large Igneous Provinces are mainly mafic magmatic provinces (having generally subordinate silicic and ultramafic components, whereas some are dominantly silicic) with areal extents >0.1 Mkm 2, igneous volumes >0.1Mkm3 and maximum lifespans of ~50 Myrs that are emplaced in an intraplate setting and characterised by igneous pulse(s) of short duration (~1-5 Myrs), during which a large proportion (>75%) of the total igneous volume has been emplaced.

How does this definition dovetail with possible LIPs on Venus? Venus surface composition, while not uniquely constrained, is widely accepted as basaltic. Areal extent is similarly easily met as tectonomagmatic volcanic rises and crustal plateaus cover areas up to 1500-2500 km diameter, whereas large coronae and radial dike swarms can exceed 350 km in diameter. Volume, while not directly constrained, does not present a serious concern given that the areal criteria of LIP is met.  The issue of time and requisite fast emplacement of <50 Myrs is not and cannot (currently) be constrained on Venus.  We simply have no means to constrain absolute time at this point in planetary exploration.  Furthermore, in many ways the near pristine character of features on Venus serves as a curse (or at least as a geological tease) toward the goal of constraining absolute time. Let me explain. Geologists commonly assume that a lack of weathering or degradation (erosion or burial) reflects relative youth; therefore, a planet surface characterized by pristine features might be assumed to be consistently young, and therefore record processes that occur quite quickly. However, interpretations of relative youth are only valid if processes that contribute to surface/feature degradation are operative at relatively short time scales.  While this logic makes sense for Earth and Mars, it is not applicable to Venus—a planet that currently lacks a hydrological cycle. Thus, it is critical to accept that we currently have no means by which to constrain absolute time on Venus; and to also accept that the incredible preservation of surface details does not reflect absolute time or age.

Because rapid emplacement is critical to any definition of LIPs, it behooves one to understand the limits of constraining time on Venus. Currently, the density of impact craters provide the only means to constrain absolute time (used quite loosely given the errors) on planet surfaces, other than Earth. For Venus, this represents a particular challenge because Venus has relatively few impact craters (~970), and lacks small impact craters (<1 km diameter). Although Venus’ impact craters represent the most pristine impact craters in the solar system their low number taken together with a near random spatial distribution prohibits robust temporal constraints for individual geomorphic features or geologic units (Campbell, 1999). Impact crater density analysis results, at best, in determination of an average model surface age (AMSA), that is, the integrated (model) age of a surface under consideration.  Whether that surface in turn represents a single geologic unit, or feature, and as such a particular AMSA represent the time of a geologic event, is a separate question that must be addressed independent of crater density. Venus records a global AMSA of ~750 +350/-400 Ma, based on total impact craters and impactor flux (McKinnon et al., 1997). Because this value is a global AMSA it represents the average (integrated) model age of Venus’ entire surface. The interpretation of this value is completely non-unique, similar in concept to an average mantle model age (e.g., Farmer and Depalo, 1983); this single global AMSA could, of course, be accommodated with a wide range of possible global surface histories (see Hansen and Young, 2007).

Impact crater dating is ultimately a statistical exercise, and includes several geological challenges. Small craters typically comprise the largest number of crater on a planet; however, Venus lacks small craters due to atmospheric screening, thus greatly limiting the resolution of a crater density surface age to a single global AMSA (McKinnon et al., 1997).  But could Venus preserve smaller domains with distinct AMSA provinces? Individual AMSA provinces must be statistically robust (a function of both local crater density and total crater population (Phillips 1993; Hauck et al, 1998; Campbell 1999)). The minimum size area that can be dated statistically on Venus by crater density alone is 20x106km2, or ~5 percent of the planet surface (Phillips et al., 1992)—or greater than two orders of magnitude larger than the minimum size area required of LIPs. Furthermore, this size area requires assumptions with regard to surface formation that severely limit the uniqueness of any temporal interpretation (Campbell, 1999). Some workers have attempted to date geological units by combining morphologically similar units into large composite regions for crater density dating (e.g., Namiki and Solomon, 1994; Price and Suppe, 1994; Price et al., 1996). However these studies lack statistical validity (Campbell, 1999), and they require the implicit (circular) assumption that similar appearing units formed at the same time, even if spatially separated.

No statistically distinct areas of crater density occur across Venus based on impact crater density alone. Therefore, any subdivision of the global AMSA, and resulting determination of robust individual AMSA provinces, requires geological criteria in addition to statistical data (crater density). Impact crater morphology might provide such criteria. Impact craters preserve morphological characteristics that record a temporal sequence of degradation in which crater halos are lost and crater troughs and interiors become progressively filled with radar-smooth material over time (Fig. 1) (Izenberg et al., 1994). Impact crater density taken together with impact crater morphology allows delineation of three separate AMSA provinces, young, intermediate, and old (Phillips and Izenberg, 1995) shown in Fig. 1. These AMSA provinces do not, however, constrain the age of individual geologic features or units; like the global AMSA, they represent AMSA’s that reflect an integrated history across these surfaces. The recorded surface histories specifically reflect only those processes that lead to formation, modification, or destruction of impact craters. It is important to realize that although the three provinces—old, intermediate and young—subdivide the global AMSA, they represent relative, not absolute, age provinces. Although no individual geologic units or features can be robustly constrained in time, individual features, or groups of features, might show spatial patterns with respect to the three AMSA provinces that might provides clues to feature formation.  However, this means that fundamentally we currently now have the means to robustly state the rate of emplacement/evolution/development of possible LIPs on Venus.

Possible LIPs on Venus

Venus displays several types of features that might represent LIPs, and which quite likely reflect different evolutionary processes: volcanic rises, large coronae, radial dike swarms, and crustal plateaus (Hansen, 2007a). Discussion here focuses on crustal plateaus, although I briefly discuss the other features.

Volcanic Rises

Volcanic rises—marked by broad domical regions (~1300-2300 km diameter;  ~1-3 km high), local radial volcanic flows, and deep apparent depths of compensation (ADC; interpreted as evidence of thermal support within the mantle)—are widely accepted as contemporary surface expressions of deep mantle plumes on thick (i.e., ~100 km) Venusian lithosphere (e.g., Phillips and Hansen, 1994; and Smrekar et al., 1997; and citations therein). A plume interpretation for volcanic rises is notable in its lack of disagreement and general consensus across a scientific community otherwise rich in diverse opinions.  The consensus no doubt results, at least in part, from the different data sets that contribute to a plume interpretation, including both geological and geophysical relations such as: size, topographic form, gravity-topography ratios, deep ADCs, and wide spread evidence of volcanic flows across their surfaces. 

Volcanic rises are divisible into three groups: rift-dominated (Beta, Atla), volcano dominated (Dione, Western Eistla, Bell, Imdr), and coronae-dominated (Central Eistla, Eastern Eistla, Themis) (Stofan et al., 1995; Smrekar et al., 1997; Smrekar and Stofan, 1999). Rift-dominated rises show the highest swells (2.5 and 2.1 km) whereas coronae-dominated rises generally show the lowest swells (1-1.5 km). Rift- and volcano-dominate rises show the highest ADCs ~175-225 km and 125-260 km, respectively, as compared to coronae-dominated rises with lower ADCs (65-120 km) (Schubert et al., 1994; Simons et al., 1997).

As a group, volcanic rises represent broadly uplifted topographic domes; they display clear evidence of extensive associated magmatism, local rifting or fracture zones, and large ADCs. All of these characteristics taken together are consistent with the interpretation that they represent surface manifestations of contemporary mantle plumes.  These features cannot be associated with plate tectonic processes, nor is there any evidence that they represent thermal insulation beneath large-tracts of continental-crust, which do not exist on Venus. Although the characteristics do not provide indisputable proof of mantle plumes, equally clearly a plume interpretation is certainly reasonable.  Although Hamilton (2005) calls for every circular and/or volcanic feature on Venus to represent a bolide impact, he does not outline any testable hypothesis for how volcanic rises could be accommodated within the context of such a hypothesis. Additionally, there is no evidence for or against a heterogeneous mantle on Venus (e.g., Meibom and Anderson, 2004; Anderson and Nathlan, 2005), and as such it seems most prudent to consider a homogeneous mantle as a first order assumption until the specific nature of proposed mantle heterogeneity is outlined within the context of a specific hypothesis or data set. Fundamentally, calling for the formation of volcanic rises by widespread melting of a heterogeneous mantle (and the inherent assumption of a heterogeneous mantle) seem equally untestable (if not more so), than a plume hypothesis. Perhaps the most interesting observation is that volcanic rises reside between 45N and 45S latitudes—reminiscent of Courtillot et al. (2003).

Large Coronae

Approximately 515 coronae, large quasi-circular tectonomagmatic features, reside on Venus; coronae range in size from 60-2600 km, but there is a strong bias toward the lower values with a median diameter of ~220 km (Stofan et al., 1992).  Coronae dominantly form chains, but also occur as clusters associated with volcanic rises, and, most rarely, as isolated lowland features. Coronae, meaning crown, was initially a descriptive term, but it has evolved into a term which commonly carries genetic connotations. Coronae are widely accepted as representing the surface manifestation of mantle diapirs (e.g., Stofan et al., 1992, 1997; Squyres et al., 1992; Janes et al., 1992). However, features collectively referred to as coronae show a wide range of topographic, structural, and volcanic characteristics, and therefore might represent more than one group of genetically unrelated features. For example, coronae could variably represent impact craters, volcanic calderas, or Rayleigh–Taylor instabilities in a density-stratified lithosphere (e.g., Squyres et al., 1992; Nikolayeva, 1993; Hamilton 1993, 2005; Schultz, 1993; McDaniel and Hansen, 2005; Vita-Finzi et al., 2005; Hamilton, 2005; Hoogenboom and Houseman, 2006).  The spatial association of corona chains and clusters with troughs and rises, respectively, favors endogenic (over exogenic) formation for these coronae. Artemis ‘Corona’ (2600 km) is an order of magnitude larger than median coronae, and over twice the diameter of the next largest corona, Heng-O (1060 km). Quetzalpetlatal (780 km), and Atahensik (~700 km; ‘Latona’ in early publications) complete the family of large coronae. Although a few other features with diameters > 600 km are included in some databases (Stofan et al. 1992, 2001), these features are: a) highly non-circular, b) display two foci, or c) represent the spatial intersection of genetically unrelated features.

Hansen (2007a) discusses the four large coronae mentioned herein, and their possible characterization as Venusian LIPs. Artemis (Fig. 2a) likely represents the surface manifestation of a plume on thin lithosphere; Heng-O’s (Fig. 2b) evolution is unclear and could represent either endogenic or exogenic driven processes. Quetzalpetlatal (Fig. 2c), requires more mapping. Preliminary work favors Quetzalpetlatal’s formation above a plume, however Quetzalpetlatal may record several distinct events, and as such, not reflect the history of a singular genetic process. Atahensik (Fig. 2d), the only large coronae to form part of a coronae chain likely represents a large diapiric structure, though the nature of the diapiric buoyancy (compositional versus thermal) is unconstrained (Hansen, 2003).

Currently Artemis has received the most attention of each of these large features with respect to geologic mapping (e.g., Brown and Grimm, 1995, 1996; Spencer, 2001; Hansen, 2002; Bannister and Hansen, 2007). Four different hypotheses are proposed for the formation of Artemis: a, b) two related to a subduction interpretation (McKenzie et al., 1992; Schubert and Sandwell, 1995; Brown and Grimm, 1995, 1996; Spencer, 2001), c) as Venus’ largest impact structure (Hamilton, 2005), and d) as the surface expression of a large mantle plume on thin lithosphere (Griffiths and Campbell, 1991; Smrekar and Stofan, 1997; Hansen, 2002, 2007b; Bannister and Hansen, 2007).

Radial Dike Swarms

Radial fracture patterns form giant radial dike swarms across Venus that could represent LIPs (e.g., Grosfils and Head 1994; Ernst et al., 2001, 2003). Some overlap exists between features mapped as radial coronae and as giant radial dike swarms; it may be the two types of features are genetically related, or the descriptive overlap may be serendipitous.  Radial dike swarms typically have radii that far exceed that of coronae annuli. For example the radial dike swarm that centers on Heng-O corona (1010 km diameter annulus) has a radius of >1000 km. The formation of giant radial dike swarms requires huge expanses of strong lithosphere; therefore the occurrence of giant radial dike swarms might place a minimum temporal limit on the existence of global scale strong lithosphere, and provide relative temporal constraints on lithosphere rheology. Indeed Venus’ giant radial dike swarms cross cut, and are therefore younger than, ribbon tessera terrain, which characterizes crustal plateaus (Ernst et al., 2003), discussed below.

Crustal Plateaus

Crustal plateaus (Fig. 3) are similar to volcanic rises in term of planform shape and size, but they differ in topographic signature and geologic evolution.  In general crustal plateaus lack rifts, large volcanoes, and coronae. Crustal plateaus host distinctive deformation fabrics (Fig. 4), ribbon-tessera terrain (Hansen and Willis, 1996, 1998). Scientists generally agree that thickened crust supports crustal plateaus, as evidenced by small gravity anomalies, low gravity to topography ratios, shallow ADCs, and consistent admittance spectra (see citations in Phillips and Hansen (1994) and Hansen et al. (1997)).  Spatial correlation of plateau topography and tectonic fabrics strongly suggests that the thickening (uplift) mechanism and surface deformation are genetically related (Bindschadler et al., 1992a, b; Bindschadler, 1995; Hansen and Willis, 1996; Ghent and Hansen, 1999).   Researchers also widely accept that arcuate-shaped inliers of characteristic ribbon-tessera terrain that outcrop across expanses of Venus’ lowland represent remnants of collapsed crustal plateaus (e.g., Bindschadler et al., 1992b, Phillips and Hansen, 1994; Bindschadler, 1995; Ivanov and Head, 1996; Hansen et al., 1997; Hansen and Willis, 1998; Ghent and Tibuleac, 2002).

Two basic questions emerge with respect to plateau formation.  1. How were plateau surfaces deformed and concurrently uplifted? And 2. How did plateaus collapse? Initially two end-member hypotheses emerged in response to the first question—the downwelling and plume hypotheses. The downwelling hypothesis involves concurrent crustal thickening and surface deformation due to subsolidus flow and horizontal lithospheric accretion associated with a cold mantle diapir beneath ancient thin lithosphere (e.g., Bindschadler and Parmentier, 1990; Bindschadler et al., 1992a, b; Bindschadler, 1995). Following the downwelling hypothesis crustal plateaus would not be considered LIPs as plateaus would not represent magmatic provinces. The plume hypothesis accommodates thickening and deformation via magmatic under-plating and vertical accretion (including volcanism) due to interaction of a large deep-rooted mantle plume with ancient thin lithosphere (Hansen et al., 1997; Hansen and Willis, 1998; Phillips and Hansen, 1998; Hansen et al., 1999, 2000). Following the plume hypothesis, plateaus could represent LIPs.

Both the downwelling and plume hypotheses call for time-transgressive deformation of ancient thin lithosphere above individual spatially localized regions, and both embrace the suggestions that a root of thickened crust supports each plateau and plateau collapse results from lower crustal flow. However, finite element modeling illustrates that the range of preserved crustal plateau morphologies and arcuate ribbon-tessera terrain inliers is difficult to achieve through lower crustal flow at geologically reasonable time scales (Nunes et al., 2004). Thus, neither elevation ranges of various plateaus nor the occurrence of ribbon tessera-terrain inliers are addressed by either the downwelling or the plume hypotheses.

Additionally, neither hypothesis addresses all characteristics of crustal plateaus, and each carries specific burdens. Challenges for downwelling include: a) mismatch of a predicted domical form (Bindschadler and Parmentier, 1990) with observed plateau shape; b) crustal thickening via lower crustal flow requires 1-4 billion years, well outside reasonable time constraints (Kidder and Phillips, 1996); and c) formation of documented short-wavelength extensional structures (ribbon fabrics) requires a high geothermal gradient (Hansen and Willis, 1998; Gilmore et al., 1998), which is difficult to justify in a relatively cold downwelling environment (Hansen et al., 1999). The plume hypothesis can accommodates a plateau shape and extensional features; however, both extensive contractional strain, and the formation of short-wavelength folds are difficult to accommodate (Ghent at al., 2005). Although the plume hypothesis addresses formation of late long-wavelength folds (or warps), which record very little shortening (<1%), early layer shortening, and/or large amounts of layer shortening present a serious challenge for the plume hypothesis (Hansen, 2006). In addition, Gilmore et al. (1998) argue that formation of ribbon fabrics requires a geothermal gradient well above that expect within the environment of a plume-lithosphere interaction.

Despite deep divides within the crustal plateau debate, SAR image mapping on both sides leads to four, mutually agreed upon, observations: 1) plateaus host both contractional structures (folds) and extensional structures (ribbons, extensional troughs, graben), which are generally mutually orthogonal; 2) multiple suites of folds, defined by wavelength, occur; 3) multiple suites of extensional structures, defined by spacing, occur; and 4) low viscosity fluid, presumably lava, fills local to regional topographic lows.  Despite these agreements, controversy exists as to the relative timing of flooding and deformation; and until recently, the amount of shortening has been unconstrained.

Detailed SAR image mapping aimed at addressing the timing of deformation and flooding, and placing limits on shortening strain, yielded new observations and refined geologic histories for plateau surfaces, resulting in the proposal of a third hypothesis—the lava-pond hypothesis (Hansen, 2006). Geologic relations call for progressive deformation of an initially very thin layer (10’s to 100 m) developed across individual plateaus.  The layer shortened forming ductile folds and extended in an orthogonal direction forming brittle ribbon structures.  With additional shortening, earlier formed short-wavelength structures were carried piggyback on younger progressively longer-wavelength folds. Local flooding accompanied progressive deformation of the increasingly thicker surface layer. Low viscosity flood material leaked from below into local structural lows. Early flooded lows were carried piggyback on younger longer-wavelength structures. Subsurface liquid (magma) formed a sharp decrease in viscosity with depth, required by structural constraints, and served as the source of flood material.

The lava-pond hypothesis calls for progressive solidification and deformation of the surface of huge individual lava ponds, each with areal extent marked by individual plateaus.  Ribbon-tessera terrain represents lava pond ‘scum’.  Individual lava ponds resulted from massive partial melting in the shallow mantle caused by large bolide (20-30 km diameter) impact on thin lithosphere (Hansen, 2006).  Melt rose to the surface leaving behind a lens of low-density mantle residuum (e.g., Jordon, 1975, 1978). This hypothesis follows the recent suggestion that the terrestrial greater Ontong-Java Plateau formed as a result of large bolide impact on thin lithosphere (i.e., Ingle and Coffin, 2004; Jones et al., 2005), following earlier suggestions (Rogers, 1982; Price, 2001).  Isostatic adjustment in the mantle, resulting from the low-density residuum lens, raised a solidified lava pond to plateau stature. Later, local mantle convection patterns could variably strip away the low-density residuum root, resulting in subsidence and/or ultimate collapse of individual plateaus. Remnants of distinctive ribbon-tessera terrain fabrics could survive as a record of an ancient lava pond. Thin surface deposits could partially or completely cover the fabrics, obscuring or erasing, respectively, evidence for individual lava ponds.

Massive partial melting within the shallow mantle could result from: a) a large bolide impact with ancient thin lithosphere, b) rise of an extremely hot deep mantle plume beneath ancient thin lithosphere, or c) a plume spawned by large bolide impact on thin lithosphere. In any case, crustal plateaus require thin lithosphere (as with the downwelling and plume hypotheses), and they owe their topographic stature to a low-density mantle residuum lens, rather than thickened crust. A bolide impact mechanism for melt-generation is favored because the formation of a lava-pond necessitates a large volume of magma at the surface at one time (and as such requires places requisite time constraints on ribbon tessera terrain formation).  Balancing formation of massive melt, yet preserving a local lithosphere able to support a huge lava pond seems a challenge to address within the context of a plume hypothesis. In contrast, a 20-30 km bolide could penetrate the lithosphere and travel into the mantle; although lithospheric penetration would damage the ‘local’ lithosphere, the lithosphere across a several thousand-km scale could retain its strength—although it might likely be riddled with fractures (Jones et al., 2005).

Although some workers state that large bolide impact cannot generate huge volumes of melt (e.g., Ivanov and Melosh, 2003), others present convincing counter arguments, particularly if a large bolide impacts thin lithosphere (Jones et al., 2005; Elkins-Tanton and Hager, 2005). Clearly such lines of inquiry are in nascent stages of investigation. Thin hot lithosphere, critical to formation of huge volumes of melt as outlined by Jones et al. (2005), might be easily accommodated on ancient Venus—widely accepted to have had a globally thin lithosphere (e.g., Solomon, 1993; Grimm, 1994; Solomatov and Moresi, 1996; Schubert et al., 1997; Phillips and Hansen, 1998; Brown and Grimm, 1999).  According to Jones et al. (2005) for impact generated melting on Earth (using the case of the Ontong Java Plateau), melt would be distributed predominantly as a giant sub-horizontal disc with a diameter in excess of 1000 km down to >150 km depth in the upper mantle within ~10 minutes, although most of the initial melt is shallower than ~100 km. The total volume of mostly ultramafic melt, would be ~3 x 106 km3 ranging from superheated liquid (100% melt, >500 degrees above solidus) within 100 km of ground zero, to varying degrees of non-equilibrium partial melt with depth and distance. These workers suggest that it would take tens of thousands of years for the melt to solidify on Earth, and it would presumably take even longer on Venus, given its dense CO2 atmosphere. Venus’ atmosphere, ~95 bars of supercritical CO2, acts more like a conductive layer, than a convection layer in terms of heat transfer, and as such serves to insulate, rather than quickly cool, the crust (Snyder, 2002). It may be that the formation and preservation of the pond scum fabric (ribbon tessera-terrain) is a function of heat-transfer properties of Venus’ atmosphere.  As noted by Jones et al. (2005), bolide impact could also spawn a thermal anomaly in the mantle, which could contribute to the thermal evolution from below, while Venus’ dense atmosphere contributed to a unique crystallization environment above.  In either case, melt would presumably rise to the surface along fractures, and would spread out, presumably as a function of surface topography. The range in individual crustal plateau planform and the specific patterns of ribbon-tessera terrain, to be determined through future mapping, might provide clues to first-order drivers of melt distribution and/or pond convection patterns

Although there is a relatively strong spatial correlation of tessera terrain fabric with crustal plateau topography, detailed mapping indicates that in some places coherent tessera fabrics extend into adjacent lowlands, as is the case of Tellus and Ovda regiones. In the context of lava pond hypothesis, ribbon terrain fabric forms on the surface, whereas the residuum—responsible for elevation—forms at depth in the shallow ductile mantle. Although we might expect broad correlation of surface and subsurface features, one can also envision situations in which the correlation would not be exact, and lava pond scum—or ribbon-tessera terrain—could be variably uplifted due to mismatch of surface lava pond and subsurface residuum by 10’s or even 100’s of kilometers. It is important to understand that the timeframe of lava pond formation and solidification would differ from isostatic adjustment due to mantle melt-residuum (the plateau-uplift event), similar to the time lag of isostatic rebound and the disappearance of continental ice sheets.  Thus a lava pond would form at the surface, solidify, and be uplifted later, and at a very different rate than lava pond solidification.

Mead Crater, Venus’ largest impact crater, at ~270 km diameter provides evidence of a bolide-impact of the postulated size (~27 km diameter) in Venus’ more recent past.  In the case of Mead Crater massive melting did not occur, presumably because the lithosphere was thick (>>10 km) at the time of impact. Thus a most critical parameter for impact-induced melting, thin lithosphere, was not met during the formation of Mead Crater, despite the presence of a large bolide impact.

The bolide impact and lava pond hypotheses also provide a mechanism to concentrate radiogenic elements in early-formed crust, with possible further differentiation into a subsurface felsic layer beneath a more mafic surface ‘scum’ (Hansen, 2007b). Crustal scale lithologic/density/radiogenic stratification at a map-scale similar to Venusian crustal plateaus is proposed for terrestrial granite-greenstone terrains (e.g., West and Mareschal, 1979; Mareschal and West 1980; Collins and Van Kranendonk, 1998; Chardon et al., 2002; Rey et al., 2003; Sandiford et al., 2004), and the bolide impact-lava pond mechanism might be applicable to Archean granite-greenstone formation (Hansen, 2007b). Surely early Earth was bombarded by bolides, which likely affected the early lithosphere. Bolides could have contributed to early mantle differentiation processes, including residuum formation, which could in turn lead to cratonization (e.g. Jordon, 1975, 1978; Bédard, 2006).


Venus, Earth’s sister planet lacks plate tectonics, yet host numerous possible examples of LIPs, which may provide clues to the formation of terrestrial LIPs. Volcanic rises represent the most likely candidates for contemporary plume signatures on thick lithosphere. Volcanic rises include three morphologically distinct groups, rift-, volcano- and coronae-dominated rises, each likely representing differences in either lithospheric properties, or relative stages of formation, or both. Four large coronae, distinguishable from other coronae based on size, are morphologically quite different from one another, and may record vastly different processes of formation. Huge radial dike swarms could represent LIPS, and might also provide critical relative temporal constraints on rheological properties of Venus’ lithosphere. Crustal plateaus, previously variably interpreted forming above deep mantle downwellings, or plumes, likely represent neither.  Their distinctive tectonic fabric with areally extensive but structurally localized lava flooding requires a massive amount of lava to have existed at one time. Crustal plateaus may represent the surface scum of huge lava ponds formed by massive partial melting in the mantle due to large bolide impact with ancient thin lithosphere. Continued study of Venus’ LIPs in combination with terrestrial LIPs will benefit our understanding of the dynamics of both planets, and terrestrial planet processes in general, and allow us to contemplate both endogenic and exogenic processes that might contribute to the formation of LIPs on terrestrial planets


Figure 1: Mollwiede projection of Magellan Venus altimetry with average model surface age (AMSA) provinces (data from Phillips and Izenberg (1995)) and major geologic features including: crustal plateaus (Alpha (pA), Fortuna (pF), eastern Ovda (pOe), western Ovda (pOe), Tellus (Te), and Thetis (pTh); large coronae, Artemis (A), Atahensik (At), Hengo (H), and Quetzalpetlatl (Q); and volcanic rises (Alta (rA), Beta (rB), Bell (rBl), Dione (rD), western, central and Eastern Eistla (rEw, rEc, rEe), Imdr (rI), Laudry (rL), Phoebe (rP), and Themis (rT)). Crater degradation stages illustrated in the cartoon show youngest (t1) to oldest (t5) changes in crater morphology; with time and degradation an impact crater loses its halo and its interior becomes radar-smooth, presumably as a result of lava fill (Izenberg et al., 1994). AMSA provinces are defined based on impact craters density and impact crater degradation stage. Old-AMSA has high impact crater density (>2.35 x 106/km3) and a deficiency in craters with halos. Young-ASA has low impact crater density (<1.85x106/km3) and a deficiency in craters with halos. Intermediate-AMSA has intermediate impact crater density and no deficiency in craters with halos (Phillips and Izenberg, 1995).  Figure modified from Hansen and Young (2006).

Figure 2: Inverted SAR (synthetic aperture radar) images of large coronae: Artemis (a), Heng-O (b), Quetzalpetlatl (c), and Atahensik (d). White regions indicate data gaps.

Figure 3: Inverted SAR image of crustal plateau Alpha Regio, a typical crustal plateau with distinctive radar-rough terrain residing in a elevated plateau above the adjacent radar-smooth lowlands; the circular feature that overlaps Alpha along its southwest margin is a younger corona.

Figure 4: Normal SAR image (radar-smooth areas are dark; radar-rough areas are bright) and geologic map of ribbon tessera-terrain, eastern Ovda Regio (from Hansen, 2006). The area encompasses the crest (~6055 km) and trough (~6054.2 km) of a prominent open fold, or warp.  Medium- and short-wavelength folds trend parallel to the long-wavelength fold crest. Orthogonal ribbon fabrics locally disrupt short- to medium-wavelength folds. Ribbon-parallel complex graben cut long-wavelength folds. All folds trend broadly parallel to one another, whereas extensional structures (ribbons and complex graben) parallel one another, and consistently trend orthogonal to fold trends. Locally, fold troughs of both short- and medium-wavelength folds are radar smooth—the result of local flooding by low-backscatter, low viscosity material. NW-trending extensional structures post-date the formation of the shortest-wavelength folds, locally cutting fold crests and troughs, including locally flooded troughs near the center of the image (black arrow). In the southeast corner of the image, periodically spaced ribbon troughs are locally filled by low-backscatter flood material (black arrows), yet other extensional structures also cut trough fill (white arrows) indicating that flooding locally both predated and post-dated formation of various extensional structures.

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