Volcanic Processes as
Alternative Mechanisms of Landform Development at
Two Candidate Lake Sites on Mars

David Leverington
Associate Professor

1.0 Overview

The past presence of lakes on Mars has been inferred from basins (e.g., impact craters) associated with inlet and outlet channels, and by the presence of features such as terraces and fan-like deposits within basin interiors [e.g., Goldspiel and Squyres, 1991; De Hon, 1992; Forsythe and Zimbelman, 1995; Grin and Cabrol, 1997; Cabrol et al., 1996, 1998, 1999, 2001; Cabrol and Grin, 1999, 2001; Ori et al., 2000; Malin and Edgett, 2003; Moore et al., 2003; Bhattacharya et al., 2005; Fassett and Head, 2005]. The candidate sites of ancient lakes have been used as a basis for theories of past Martian climatic conditions, and have been suggested as possible locations for future astrobiological investigations [e.g., Goldspiel and Squyres, 1991; De Hon, 1992; Newsom et al., 1996; Ori et al., 2000; Cabrol and Grin, 2001; Rathbun and Squyres, 2002; Murray et al., 2005].

Recent work has suggested that numerous hypothesized Martian lake and channel features of Hesperian and Amazonian age, rather than having formed in association with lacustrine and fluvial environments, are likely to have instead formed by lava flow related to regional volcanic resurfacing [e.g., Leverington and Maxwell, 2004; Greeley et al,. 2005; Leverington, 2006]. The basic concepts discussed in two of these studies [Leverington and Maxwell, 2004; Leverington, 2006] are outlined below. Additional general information is available here.

The material below was compiled and published prior to 2007, and therefore does not cover more recent findings of relevance [e.g., Erkeling et al., 2011].

2.0 A Channel System in the Memnonia Region of Mars

A remarkable crater and channel system is centered in a ~250 x 350 km region located in western Memnonia, an area of heavily cratered Noachian-aged highlands [Mutch and Morris, 1979; Scott and Tanka, 1986]. The primary crater of interest in this system (referred to here as the “central crater”) is centered at ~174.8o W and 14.6o S, and has a diameter of 45 km. This crater has a maximum rim elevation of ~1650 m and a minimum floor elevation of ~375 m. The crater has an almost continuous inner terrace that is tilted toward crater center with convex form [e.g., Forsythe and Zimbelman, 1995; Ori et al., 2000], with maximum radial widths of ~5 km. In places, discontinuous segments of additional terrace-like features are also present. Wrinkle ridges and lobate-margined flow units are associated with materials that have accumulated inside the crater; these features, as well as moat-like features peripheral to scarps formed by units of interior crater fill, are also present in the fill materials of numerous nearby craters. The materials that comprise exposed fill units of the main crater (including the surfaces of the terrace and of the inner basin) preserve a common surface texture and a cratering record that differs from that of adjacent highlands for craters with diameters of less than ~300. Highland features in this region display a softened appearance relative to infill materials.

The inlet channel of the central crater is sinuous and flat-floored. At its mouth, the channel has a width of ~3.5 km and a floor elevation of ~620 m. The main inner terrace of the crater extends more than 10 km up the inlet valley [e.g., Forsythe and Zimbelman, 1995; Ori et al., 2000]. To the south, the inlet channel has a width of ~800 m, and forms part of a low-order system of relatively narrow and low-relief channels (with typical widths of ~800 m and local depths of ~10-25 m) that, together with the craters to which some of these channels are connected, extends throughout much of the inlet basin. The terrain of the inlet basin is mostly rounded and subdued, with the steepest and most dissected slopes mainly confined to relatively small areas at higher elevations.

The outlet channel of the central crater extends northeastward into a 4 to 6 km wide valley that terminates at the highland-lowland boundary [Scott and Tanaka, 1986], ~160 km from the northern perimeter of the main terraced crater, and immediately south of deposits of the Medusae Fossae Formation [e.g., Bradley et al., 2002]. The channel begins at a blunt amphitheater-like headwall at an elevation of ~570 m and with a maximum channel width of ~1.8 km. The channel sinuously extends northeastward with a gradual decrease in depth corresponding to a reduction in local slope, fading into relatively flat-floored valley fill about 40 km from the channel head. The channel’s deepest elevation relative to its rim is ~325 m. A distal segment of the channel appears further down the valley where slope increases, widening to ~3 km before narrowing and fading into deposits of the northern lowlands. The outlet valley has one major tributary valley that intersects it from the southeast, about 50 km from the mouth of the outlet channel. Crater ejecta covers part of the valley floor near this intersection. The full length of the main outlet valley is ~190 km. The gradient of the outlet channel ranges between ~6 and 18 m / km, and the overall gradient of the outlet valley is ~12 m / km.

Figure 1: Map of MOLA elevation data superimposed on shaded relief of the study area, showing the locations of the main terraced crater a, the inlet basin b, and the outlet valley c [Leverington and Maxwell, 2004]. Contour interval is 250 m. The locations of Figures 2, 6, and 7 are indicated. Topographic data after MEG128 model of Smith [2003].

Figure 2: The large terraced crater at the center of the study region (crater diameter = 45 km) [Leverington and Maxwell, 2004]. Features are labeled as follows: a: inlet; b: terrace; c: outlet. The locations of Figures 3a, 3b, 5, and 8 are indicated. Elevation profile across line x-y is shown at bottom [after Smith, 2003]; note the prominent terrace along the inner crater flanks. Image location given in Figure 1. Viking Orbiter frame 438S12.

Figure 3: Floor materials of the terraced crater showing: a: wrinkle ridges, and b: lobate margins [Leverington and Maxwell, 2004]. Floor materials with similar features are found inside many craters and topographic depressions in the region. Image locations are given in Figure 2. Themis frame V04688002.

Figure 4: Terraces (t) and moat-like features (m) that are peripheral to fill deposits in craters of the inlet basin (also see Figure 6) [Leverington and Maxwell, 2004]. The peripheral moats are exterior to and partly formed by narrow adjacent rises. Viking Orbiter frames: a, 437S14; b, 437S20; c, 437S19; d, 437S14; e, 436S17; f, 437S14.

Figure 5: Surface texture of floor materials of the terraced crater: a: crater rim; b: terrace; c: bottom of terrace scarp; d: crater floor [Leverington and Maxwell, 2004]. Image location is given in Figure 2. Mars Orbiter Camera frame M0903914.

Figure 6: Inlet channel of the main terraced crater, and terraced craters associated with subtle channel features that feed into the inlet channel [Leverington and Maxwell, 2004]. Features are labeled as follows: a: elongate terraced crater with peripheral moat; b: circular terraced crater; c: channel extending from crater a; d: inlet area of main terraced crater; e: breached crater. Image location is given in Figure 1. Viking Orbiter frames 437S14, 437S15, and 437S16.

Figure 7: Outlet valley of main terraced crater [Leverington and Maxwell, 2004]. Features are labeled as follows: a: outlet head; b: outlet valley floor; c: tributary valley; d: distal outlet channel. Image location is given in Figure 1. Viking Orbiter frames 599A70 and 599A72.

Figure 8: Head of outlet channel [Leverington and Maxwell, 2004]. Image location is given in Figure 2. Mars Orbiter Camera frame M0702911.

2.1 Are There Potential Issues with Lacustrine Interpretations of These Landforms?

Many workers have suggested that the past existence of lacustrine and fluvial environments on Mars might account for the association of channels, inner terraces, and fan-like features with topographic basins (many of which are impact craters) [e.g., Forsythe and Zimbelman, 1995; Cabrol and Grin, 1999; Ori et al., 2000; Cabrol et al., 2001]. However, recognized modes of formation of lacustrine terraces and channels cannot easily account for the nature of the candidate crater-lake system examined here. For example, extensive lateral dimensions are typically required of water bodies in order to develop the fetch necessary for significant wave-cutting action. It is unlikely that a maximum fetch of only 45 km in the central crater could have allowed wave action to form terraces that are orders of magnitude larger than most terrestrial lacustrine counterparts [e.g., Gilbert, 1885, 1890; Brophy, 1967; Matmon et al., 2003]; attenuation of wave energy by a growing terrace and shelf should have further inhibited widening of these features [see, e.g., Trenhaile, 1983, 2000, 2002]. The terrace of the central crater is large relative even to most terrestrial marine terraces [e.g., Keraudren and Sorel., 1987; Massari et al., 1999; Polenz and Kelsey, 1999; Tortorici et al., 2003; Maeda et al., 2004; Yamaguchi and Ota, 2004].

Both the inlet and outlet channels possess properties that are not clearly supportive of aqueous origins involving association with an ancient crater lake. For example, the rim of the outlet channel and the fill of the outlet valley are not fluvially dissected, even though such modification would be expected under the same environmental conditions that would have maintained a crater lake. Formation of the outlet channel by groundwater sapping is not necessarily compatible with active formation of channels of the inlet basin by substantial water flow, and with ponding of these waters to form a crater lake drained by the outlet channel.

Aeolian, erosional, and structural models for the origin of the terrace of the central crater are, as with the lacustrine model, problematic. For example, formation of the terrace by aeolian infilling followed by selective aeolian exhumation of materials in the crater center is not consistent with the lack of similar deposits in numerous nearby craters, nor would such exhumation necessarily be expected to produce a roughly continuous and symmetric terrace that also continues into the crater’s inlet. Formation of the terrace following impact or through subsequent crater erosion is not likely, given the absence of obvious evidence for listric-normal slope failure [e.g., Ori et al., 2000]; the extension of the terrace into the inlet valley also weakens this hypothesis, in that structural or mass-wasting features might be expected to be related to wall failure, and oriented roughly symmetrically about the center of the crater.

2.2 Do Viable Alternatives Exist to Aqueous Interpretations?

The nature of the central crater, the terraces and channels of the inlet basin, and the channel of the outlet valley are not strongly suggestive of formation and evolution by fluvial and lacustrine processes. However, the morphologies of and interrelations between these features are consistent with formation by extrusive igneous processes, with interior crater fill having accumulated through deposition and subsidence of materials erupted at the surface during effusive volcanic activity maintained by a system of intrusions, and with channels having formed as volcanic rilles that transported lava from overflowing sources and other filled depressions.

Evidence in support of a possible volcanic origin for the system under study is grouped here into three categories: 1) crater fill and associated terraces; 2) outlet channel; and 3) inlet basin.

2.2.1 Crater Fill and Associated Terraces

The interior units of many of the craters of the study region are associated with: a) wrinkle ridges that are suggestive of subsidence and lateral compression of layered volcanic units; b) surface units with lobate margins and peripheral moats that are consistent with formation through flow and subsidence of volcanic materials; c) textural, crater-preservation, and thermal characteristics that are collectively consistent with volcanic materials that are relatively dense and consolidated; and d) inner crater terraces that are consistent with subsidence or drainage of volcanic materials.

Wrinkle ridges are present in the fill of the central crater as well as that of craters of the inlet basin. These features have widths of ~1-2 km, lengths that in some cases exceed 10 km, and heights of meters to tens of meters; these dimensions fall within the size range of small Martian wrinkle ridges [e.g., Watters, 1988; Schultz, 2000; Mont(si and Zuber, 2003]. Wrinkle ridges are sinuous and elongate topographic highs that are found on Mercury, the Moon, Venus, and Mars, and are believed to form through horizontal shortening of near-surface layered deposits [e.g., Watters, 1988, 1991; Golombek et al., 1991; McGill, 1993; Schultz, 2000]. Although the occurrence of wrinkle ridges on the terrestrial planets is not in principle restricted to materials with a volcanic origin [see, e.g., Schultz, 2000], most or all clear examples of these features are found in materials interpreted to be layered and volcanic in origin [Watters, 1988]. On this basis, and on the basis that wrinkle ridges are very commonly formed in lunar and martian basaltic deposits [e.g., Muehlberger, 1974; Maxwell, 1978; Solomon and Head, 1979; Watters, 1988; Head et al., 2002], the existence of wrinkle ridges in the study region is interpreted as suggestive of a volcanic origin and layered nature for materials that have accumulated in topographic depressions.

Many geological units deposited in craters and other depressions in the study region have lobate margins that typically have local elevations of <10 m, but can reach ~10-30 m in height. The thicknesses and appearance of these units are comparable to, for example, volcanic flows of the Martian highlands and lowlands [e.g., Schaber, 1980; Baloga et al., 2003; Ivanov and Head, 2003] and prominent lunar volcanic flows in Oceanus Procellarum and Mare Imbrium [e.g., Schaber, 1973; Schultz, 1976; Zimbelman, 1998; Hiesinger et al., 2002]. Some lobate units in the study region form prominent scarps that are peripheral to crater-fill deposits, and that can resemble the peripheral flow features of lunar lava lakes [e.g., El-Baz and Roosa, 1972; Schaber et al., 1976] and lunar craters containing subsided accumulations of volcanic materials [Schultz, 1976].

Exposures of the terrace and inner fill materials of the main crater have an identical surface texture as viewed in high-resolution MOC images. This texture is similar to that of materials that have accumulated on the flanks of martian volcanoes such as Hadriaca Patera, and is consistent with common volcanic origins. The nighttime thermal properties of crater fill of the central terraced crater and numerous other craters in the study region are suggestive of surface materials that are relatively dense and consolidated compared with materials in surrounding areas (e.g., Themis frame I06841006).

Terraces that may not have formed through mass-wasting or tectonic processes are found on the inner flanks of numerous craters in the study region, and include the relatively narrow terraces of craters in the inlet basin and the relatively broad terrace associated with the central crater. The formation of terraces and moats inside impact craters and along the perimeters of fill accumulations through subsidence of volcanic materials is believed to have occurred on the Moon [e.g., Holcolm, 1971; El-Baz, 1972; Greeley, 1976; Schultz, 1976; Young, 1976; Greeley and Spudis, 1978], perhaps involving such well-known terrestrial processes as degassing or cooling of subsurface magma [see, e.g., Francis et al., 1993], or involving subsidence through loading. The lunar terraces are morphologically similar to the narrow terraces found in the study region, and are suggestive of formation through analogous igneous processes.

While the broad terrace of the central crater could have also formed through subsidence of volcanic materials, other mechanisms are possible. The cooling and crystallization of a lava lake, taking place along all exterior surfaces of the lava body [e.g., R(pke and Hort, 2004], could, depending on associated degassing processes, result in terrace-forming subsidence of the cooled surface. Alternatively, there are abundant terrestrial examples in which inner terraces have formed in calderas by substantial collapse caused by evacuation of subsurface magma chambers by degassing, lateral intrusion, or drainage events [e.g., Young, 1976; Francis et al., 1993; Geshi et al., 2002; Gottsmann and Rymer, 2002]. Repeated filling and drainback of terrestrial lava lakes, driven by cyclic processes such as vesiculation within the magma column beneath the source vent, can also produce distinct terraces [e.g., Burgi et al., 2002; Barker et al., 2003]. Based on these terrestrial analogs, it is possible that the broad terrace in the main crater formed through the occurrence of at least one drainage event of a lava lake located inside the crater. If such an event took place, then the terrace in this crater would correspond to an earlier, higher filling level, and the outlet channel would have acted as a conduit for this and subsequent releases of lava. Lava terraces are present in Bowditch, a 35 km long lunar crater located northeast of Mare Australe (see figure below). These terraces are up to 3.5 km wide and 50 to 200 m high [Young, 1976], and are believed to mark the high level of a lava lake that breached the crater rim [El-Baz, 1972; Young, 1976]. A large patch of mare-like material (Lacus Solitudinis) extends ~120 km in a southeastward direction from the breach. Crater Bowditch and its terraces are of comparable size, morphology, and appearance to the central crater of the Memnonia study region. The similarity between the appearance, dimensions, and geological context of these features is consistent with their common formation by breach and outflow of pooled lava bodies.

2.2.2 Outlet Channel and Valley

The outlet channel is a sinuous single channel that heads at full width and gradually tapers in width in a downchannel direction, has sharp and undissected rims, and progressively shallows until it disappears into relatively flat-floored valley materials before re-appearing at the steeper slopes that extend onto the volcanic plains of the northern lowlands. As such, the morphological characteristics of this channel closely match those of many sinuous rilles on the Moon [e.g., Schubert et al., 1970; Greeley, 1971a; Schultz, 1976; Guest and Murray, 1976; Strain and El-Baz, 1977; Hulme, 1982]. The full length of the outlet channel, including the flat-floored section where surface expression of the channel is absent, is ~190 km; the greatest channel widths range between ~1.8 and 3 km, and maximum depth is ~325 m. These dimensions are generally consistent with those of such lunar rilles as Hadley, which is ~135 km long and averages 1.2 km in width and 370 m in depth [Greeley, 1971b].

Lunar sinuous rilles are believed to have formed in a manner analogous to terrestrial volcanic channels and lava tubes [e.g., Oberbeck et al., 1969; Greeley, 1971b; Cruikshank and Wood, 1972; El-Baz and Roosa, 1972; Young et al., 1973; Carr, 1974], which form as conduits for lava streams in effusive flows [see e.g., Stephenson et al., 1998; Calvari and Pinkerton, 1998; Kauahikaua et al., 1998]. Levees often form along active flow margins of lava channels, and formation of fully-enclosed lava tubes can result from processes such as accretion on levees or aggregation of floating crustal rafts [e.g., Peterson et al., 1994; Kauahikaua et al., 1998]; formation of crusted roofs can dramatically decrease cooling rates by insulating the core of the flow from direct radiative and convective cooling, allowing basaltic lava flows to form with lengths of hundreds of kilometers [e.g., Keszthelyi and Self, 1998; Sakimoto and Zuber, 1998; Harris and Rowland, 2001]. The roofs of lava tubes can partially or completely collapse following drainage of parent flows [e.g., Papson, 1977; Calvari and Pinkerton, 1998], producing the appearance of discontinuous channels. Steeper segments of terrestrial and lunar lava tubes typically form no or weak roofs, while segments with low slopes can form stronger roofs and are less likely to drain [e.g., Greeley, 1971b; Sakimoto and Zuber, 1998]; as a result, there can be preferential preservation of the roofs of tubes along segments of low slope, a situation that may apply to the morphology of the outlet channel in the Memnonia study region. The absence of obvious accumulations of material that flowed out of the outlet channel is a quality that is typical of lunar sinuous rilles [e.g., El-Baz et al., 1972], and in the case of the Memnonia outlet channel could be related to burial by later volcanic flows. The absence of fluvial dissection of the margins and interior of the outlet channel is consistent with the channel’s origin as a volcanic feature that formed under relatively dry climatic conditions that have persisted from the Hesperian until the present.

Could the flow of lava have breached the surface or subsurface of the northern rim of the central crater? Volcanic rilles usually form as constructional features that emplace lava flows, and incision is often not important in their formation [e.g., Young et al., 1973; Carr, 1974; Greeley, 1977]. Furthermore, aspects of the morphologies of deep lunar lava channels can in some cases simply be attributed to factors such as preexisting topography and constructional processes associated with the formation of levees [e.g., Sparks et al., 1976; Greeley, 1987; Spudis et al., 1988] (indeed, the morphological characteristics of the full length of the Memnonia outlet channel north of the crater breach are consistent with that of a constructional channel formed during flow and infill of lava in a pre-existing valley). Nevertheless, while there is much that is not understood about the environmental and eruptive conditions necessary to support the formation of erosional channels [e.g., Fagents and Greeley, 2001], thermal and mechanical processes of erosion by flowing lava have been recognized as important through both theoretical studies [Huppert et al., 1984; Huppert and Sparks, 1985; Kerr, 2001] and terrestrial field studies of active and ancient flows [e.g., Peterson and Swanson, 1974; Barnes and Barnes, 1990; Greeley et al., 1998; Kauahikaua et al., 1998, 2002, 2003; Williams et al., 2004]. A basaltic erosive rate of 0.1 m / day over a period of 60 days has been measured for a relatively small lava stream with laminar flow at Kilauea Volcano in Hawaii, suggesting a capacity for local-scale basaltic erosion to incise tens of meters into a basaltic substrate over a period of less than one year [Kauahikaua et al., 1998]. The erosive capacity of lava increases with higher temperatures and volumes, lower viscosities, and increased turbulence [see, e.g., Greeley et al., 1998; Williams et al., 1998], suggesting a far greater capacity for erosion for lunar and martian flows associated with large rilles. Capacity for erosion may be still greater for cases in which incision involves a regolith substrate rather than bedrock.

2.2.3 Inlet Basin

The surface of the inlet basin is crossed by shallow channels that connect craters and other depressions that have been filled with deposits interpreted above as volcanic in origin. The distribution and morphology of these channels are consistent with formation by overtopping of confining rims of topographic lows into which volcanic materials accumulated, perhaps in a manner analogous to the formation of channels by constricted overflow of lava lakes on the Moon [e.g., Schaber, 1973]. Channels in the inlet basin ultimately converge on the inlet of the central crater; the extension of the crater terrace into the inlet valley is suggestive of a terrace origin that post-dates formation of the crater [Ori et al., 2000], and is consistent with formation by lava flow through the inlet during periods that followed reductions from earlier, higher levels. The inlet breach in the crater wall could conceivably have formed by lava erosion, if lava ponded at the lowest part of the inlet basin, although initial formation of the inlet breach through other processes is also possible.

Lava flows have the capacity to follow routes defined by the gradients of pre-existing topography [e.g., Swanson and Wright, 1978; Reidel, 1998; Branca, 2003], to incise channels [e.g., Kauahikaua et al., 1998], and to flow hundreds of kilometers [e.g., Wilson and Head, 1994]. The study region is located within a broad area that has clearly been subjected to highland volcanism. Extensive plains units characterized by flow lobes and wrinkle ridges, interpreted as highland flows of low-viscosity lava erupted from numerous sources at high rates, are found immediately south of the divide that defines the inlet basin [Scott and Tanaka, 1986]; there are several breaks in the divide that separated these flows from the inlet basin (e.g., at 173o52’ W, 16o48’ S), providing possible locations from which the inlet basin could have been repeatedly flooded by lava. Volcanic features located within several hundred kilometers of the study region (in adjacent regions of Memnonia, Aeolis, and Phaethontis quadrangles) include the fissure- and caldera-like depressions of Scott and Tanaka [1986], faults and graben of Memnonia Fossae [e.g., Tanaka and Chapman, 1990; Wilson and Head, 2002], and the lowland shield volcano, Apollinaris Patera.

Figure 9: Examples of flow fronts in Mare Imbrium [Leverington and Maxwell, 2004]. Lunar Orbiter frames V-160-H2 and V-161-H2.

Figure 10: Moat (m) believed to have been formed through subsidence of volcanic deposits that once covered this crater [see, e.g., Schultz, 1976] [Leverington and Maxwell, 2004]. Lunar Orbiter frame III-150-M.

Figure 11: Lunar impact-crater terrace believed to have been formed through subsidence of volcanic fill [see, e.g., Schultz, 1976] [Leverington and Maxwell, 2004]. Lunar Orbiter frame IV-195-H2.

Figure 12: Terraces (marked t) located in the Flamsteed region of the Moon and believed to have formed through subsidence of volcanic deposits [see, e.g., Schultz, 1976] [Leverington and Maxwell, 2004]. Lunar Orbiter frame III-200-M.

Figure 13: Lava terraces (t) in Bowditch (a), a 35 km long crater in the highlands of the lunar farside, northeast of Mare Australe [Leverington and Maxwell, 2004]. The terraces are believed to mark the high level of a lava lake that breached the crater rim, forming a large patch of mare-like material (Lacus Solitudinis; b) [El-Baz, 1972; Young, 1976]. Apollo 15 Metric Camera frame AS15-M3-2628 (see also Apollo 15 Panoramic Camera frame 9965).

Figure 14: Comparison of (a) distal outlet channel of main terraced crater (Viking Orbiter frame 439S07) with lava conduits at (b) Ceranius Tholus, Mars, (Themis frame V01002003) and (c) Hadley Rille (Lunar Orbiter frame V-105-H2) [Leverington and Maxwell, 2004].

Figure 15: The outlet channel of lunar crater Krieger (a) (Apollo 15 Panoramic frame 0327) [Leverington and Maxwell, 2004].

Figure 16: Comparison between the morphology of channels seen in the inlet basin of the Memnonia study region (a, b, c; Viking Orbiter frames 437S14, 438S13, 437S17, respectively) and those on the flanks of martian volcanoes (d – Ceranius Tholus, e – Hecates Tholus, f – Ceranius Tholus; Themis frames V02787010, V03128003, V04310006) and volcanic rille features in the Aristarchus Plateau region of the Moon (g, h, i; Apollo 15 Panoramic frames 0326, 0326, 0328) [Leverington and Maxwell, 2004]. The forms of the Memnonia channels, which are typically low-order channels characterized by coarsely uniform widths and abrupt heads and terminations, are not distinct from the martian and lunar volcanic channels.

2.3 Discussion

The igneous hypothesis outlined above would have essentially involved past development of a Hesperian-aged interlinked system of surface lava conduits and overflowing basins, with candidate lava sources including those of the extensive region of highland flows located immediately to the south of the study region. Under this hypothesis, most or all craters and other topographic depressions in the study region would have been passive repositories for multiple layered volcanic deposits, and, at times, lava lakes; not being sustained by mass exchange with subsurface magma chambers, individual lava lakes would likely have been very short-lived [see, e.g., Worster et al., 1993; Burgi et al., 2002]. Flow of volcanic materials through the inlet basin, central crater, and outlet channel would have resulted in voluminous accumulation on the plains north of the dichotomy boundary.

Hypotheses for the origins of channels and basin units in the Memnonia study region are testable. For example, aqueous development would, for many models, be expected to have resulted in widespread aqueous alteration of involved units. The materials that accumulated in basins should also have properties consistent with deposition under lacustrine conditions, and associated channels should show some evidence for transport of water and sediment. In contrast, volcanic development could conceivably have operated under drier conditions that favored preservation of original mineralogies (generally dry conditions are expected to have prevailed globally over much of the Hesperian and Amazonian). Volcanic development would also be expected to have resulted in emplacement of lava flows in basins located along channel systems, as well as at the terminal basin of the system (i.e., in the northern lowlands, which commence just north of the dichotomy boundary in this region).

3.0 The Palos Crater / Tinto Vallis System, Northernmost Hesperia Planum, Mars

Palos crater is located at ~2o40’ S, 110o54’ E, in a region of Noachian highlands that lies immediately north of the extensive ridged plains that comprise the flanks of Tyrrhena Patera, a large low-relief volcanic rise located in the Hesperia Planum region of Mars [e.g., McCauley et al., 1972; Plescia and Saunders, 1979; Greeley and Guest, 1987; Greeley and Crown, 1990]. Impact craters with diameters larger than ~20 km are generally highly degraded in this region, and lack well-defined crater rims and ejecta blankets. The main crater of the candidate crater-lake site has a diameter of ~50 km and is also highly degraded.

The rim of the main crater is crossed in the south by an inlet channel that branches upslope into a drainage system that can be divided into western and eastern sections [e.g., Carr, 1995; Scott et al., 1995]. The eastern section consists of a simple sinuous channel, Tinto Vallis, that extends upslope toward the south. The southern segment of this channel, and the lengths of several associated tributaries, are defined by chains of rimless collapse pits that indicate that these portions of the channel system are partially roofed [e.g., Carr, 1995; Cabrol and Grin, 2001]. The width of the sinuous channel ranges mainly between ~2.5 and 3.5 km, with channel widths greatest for more northerly channel segments. The floor of the channel is roughly 600 to 700 m below the upper rim of the channel, and in places is mantled by light-toned aeolian megaripples (e.g., MOC image M0704217); typical channel width-to-depth ratios range between ~4:1 and 6:1. A large rimless collapse pit in the south has a base that is roughly 700 m below its rim; other pits in the area are more subtle and shallow, with depths mainly less than 300 m. Ultimately, the eastern section of the inlet system heads on the flanks of Tyrrhena Patera. Elevations of the channel floor, derived from minimum-elevation shot points located along selected MOLA transects of the channel, are given below for the reach between the inlet breach and a large area of collapse; the elevation of the channel floor drops ~600 m over a distance of ~180 km for this reach, for an overall gradient of ~3.3 m / km, or less than 1%. The western section of the inlet system consists of a localized but complex system of valley networks that in places is dendritic in nature. The western part of this system heads at a local topographic divide, while the channels of the southern part of this system shallow and fade into terrain to the south toward Tyrrhena Patera.

The floor of the central crater has an average elevation of approximately -690 m and a standard deviation about this value of ~60 m; the lowest areas of relatively broad extent (i.e., areas not restricted to the sub-basins of individual impact craters) approach approximately -800 m and are found in the northeastern quadrant of the crater interior. The highly degraded rim of the crater rises ~650 to 1200 m above the average floor elevation. The surface texture of the infill of the central crater is generally complex, and infill materials show evidence of layering. The topography of fill materials of the central crater is subdued, with relief less than ~100 m over distances of tens of kilometers. At relatively high spatial resolutions, the surface morphology of fill materials ranges between relatively smooth surfaces and very irregular or hummocky tracts. Much of the crater floor is densely pitted by craters with diameters less than ~100 m; the ejecta blankets of two proximal upland craters partly extend across infill materials. In places, crater-infill deposits have distinctly lobate margins that rest against the main interior crater walls. Terrace-like landforms of relatively limited extent are present in the main crater, but the crater interior lacks a distinct and continuous terrace along its periphery. Aeolian deposits, most commonly in the form of large fields of light-toned megaripples, partly superpose extensive tracts of the crater-fill materials (e.g., MOC image E0101877) as well as the surfaces of nearby uplands (e.g., MOC image M0300742). Low mounds form regions of hummocky terrain in central and eastern areas of the crater interior, and there is a partial correspondence between this terrain and relatively cool areas in nighttime thermal images.

A north breach in the rim of the main crater acted as a conduit for fluid flow from the main crater, opening toward Amenthes Planum. The elevation of plains materials at the north breach is approximately -695 m, and although elevations within ~200 km of the north breach mainly range within only ~100 m of this value, the plains of Amenthes Planum ultimately drop more than 2 km before reaching the basin of Isidis Planitia, located ~800 km to the northwest. The Hesperian wrinkle-ridged infill of the broad trough that contains Amenthes Planum [Hiller, 1979; Tanaka et al., 2005] is contiguous with the infill of the central crater. The position of the outlet breach of the central crater marks the location of an outlet channel that crosses the interior of the main crater and that has been almost completely filled by materials with a surface appearance similar to that of other surficial materials within the crater; the characteristics of surface materials of the filled outlet channel are spatially variable with regard to nighttime temperatures, suggesting that these materials may be variably consolidated, or mantled by materials of varying particle sizes or limited distribution. Extensive ridged plains are located both downslope and upslope of the main crater; the plains downslope of the crater ultimately extend to the plains of Isidis Planitia, whereas the plains upslope from the main crater form the northern flanks of Tyrrhena Patera.

Figure 17: Map of MOLA elevation data superimposed on shaded relief for the Tyrrhena Patera region, showing the locations of a regional Viking mosaic (Figure 18), a mosaic of nighttime thermal images of the candidate crater-lake site (Figure 23), and images of ridged plains (Figures 25 and 26) [Leverington, 2006]. Topographic data after Smith et al. (2003).

Figure 18: Viking Mars Digital Image Model (MDIM) of the northern flanks of Tyrrhena Patera and associated uplands, showing the location of Palos crater (Figures 19 and 20) and the locations of images of crater-floor materials (Figure 22) [Leverington, 2006]. Image location is given in Figure 17. Illumination is from the right.

Figure 19: Palos crater (A), a candidate crater-lake site located immediately north of Tyrrhena Patera (site 2.7S249.2 of Cabrol and Grin, 2001; see also Carr, 1995) (Viking frame 379S45) [Leverington, 2006]. The outlet breach (B) of the site opens to a broad plain (Amenthes Planum, ap) that ultimately extends northwest and downslope to Isidis Planitia. The inlet breach (C) branches upslope into western and eastern sections. The eastern section, Tinto Vallis, is a simple sinuous channel (D) with roofed tributaries (E) that are discontinuously expressed by rimless collapse pits (arrows); the partially-roofed system appears to originate upslope on the flanks of Tyrrhena Patera (F). The western section of the inlet system is a very complex and, in places, dendritic channel and valley system (G and H). Other impact craters in the region are associated with channel systems that also appear to ultimately extend from Tyrrhena Patera (e.g., I). Some impact craters in the region are partially covered by irregular or chaotic accumulations of material (e.g., J). Impact craters with ejecta blankets that extend partly across infill materials of the central crater are indicated (c1 and c2). Image location is given in Figure 18. Illumination is from the right.

Figure 20: MOLA topography superimposed on shaded relief of the candidate crater-lake site (database extent is given in Figure 18; topographic data after Smith et al., 2003); note that parts of the northern reach of the main inlet channel are imperfectly represented in the gridded MOLA database as a consequence of the relatively small size of the channel and the uneven coverage of original MOLA transects [Leverington, 2006]. Individual profiles across selected features are given at right (MOLA orbit numbers: a, 12297; b, 13573; c, 12800; d, 13039). Minimum elevations at profiles a, b, and c are -618 m, -213 m, and -162 m, respectively. The average elevation along profile d is approximately -670 m and the standard deviation about this average is ~17 m. Profiles contain no vertical exaggeration.

Figure 21: Plot of channel-floor elevations against distance along the most distinct reach of the partially-roofed sinuous inlet channel [Leverington, 2006]. Elevations correspond to the minimum-elevation shot points of selected MOLA transects across the channel. The channel reach represented in this plot extends from the inlet breach (C in Figure 19) upslope to a large area of collapse (E in Figure 19), corresponding to an along-channel distance of ~180 km. The lowest and highest elevations in this plot are -642 m and -33 m, respectively; the overall gradient of this reach is less than 1%. The channel appears to extend upslope beyond this particular reach along several paths (e.g., F in Figure 19). The channel is relatively narrow, and thus a substantial proportion of the scatter in the channel-floor elevations presented in this plot is expected to be an artifact of the sampling characteristics of the MOLA instrument itself; that is, with a footprint of 160 m and a 300 m separation between shot points, the lowest value in a given MOLA-transect dataset is unlikely to correspond precisely to the lowest transect elevation on the ground.

Figure 22: Floor materials of Palos crater [Leverington, 2006]. Image locations are given in Figure 18. A filled channel extends northward through the outlet breach (at “A” in figure a). Individual layers or sets of layers can be identified in some locations (e.g., at black arrows in a); some units have lobate margins in plan view (e.g., at white arrows in figures a and c). Low mounds form tracts of hummocky terrain in central and eastern areas of the crater interior (M in figures a and b). The locations of two parts of a high-resolution MOC image (Figure 23) are given in figure a. THEMIS VIS (Band 3) images V02531001 (a), V05527002 (b), and V09434001 (c). Illumination is from the left.

Figure 23: Two parts of a high-resolution MOC image of floor materials of Palos crater [Leverington, 2006]. Light-toned aeolian megaripples (e.g., at arrow in b) superpose some fill materials. Image locations are given in Figure 22. Mars Orbiter Camera frame M0401804. Illumination is from upper-left.

Figure 24: Mosaic of THEMIS nighttime infrared images (band 9) [Leverington, 2006]. Image location is given in Figure 17. THEMIS frames I07218007, I07580007, I07967012, and I07605012.

Figure 25: Ridged plains of Amenthes Planum, located southeast of Isidis Planitia [Leverington, 2006]. Image location is given in Figure 17. THEMIS daytime infrared frame I01519012 (band 9). Illumination is from bottom left.

Figure 26: Ridged plains on the distal northern flanks of Tyrrhena Patera [see also Greeley and Crown, 1990] [Leverington, 2006]. Image location is given in Figure 17. THEMIS daytime infrared frame I09746001 (band 9). Illumination is from the left.

3.1 Past Aqueous Interpretations of the Palos Crater / Tinto Vallis System

Although the candidate crater-lake site described above has not been subjected to detailed examination in the past, several studies have considered the site in general terms and concluded that its features are suggestive of formation by aqueous processes [Carr, 1995; Scott et al., 1995; Cabrol and Grin, 2001]. Carr [1995] has suggested that karst-like processes may have been involved in the formation of features of the inlet system of the candidate crater-lake site. Such a hypothesis is very attractive in that it can potentially account for much of the drainage system to the south of the site, including the partially roofed segments and tributaries of the main inlet channel, which under this scenario would have developed through subsurface erosion or solution by groundwater. Cabrol and Grin (2001) support this general hypothesis, noting that the depressions at the site are suggestive of a volatile-rich subsurface (although their analysis erroneously relates the subsurface conduits of the site to Shalbatana Vallis, located in Xanthe Terra).

None of the fundamental characteristics of the Palos crater site appear to preclude previously-hypothesized aqueous origins for channels and crater fill. Indeed, an aqueous mechanism of formation of the key features of the candidate crater-lake system (including the inlet channel system and the infill of the main crater) would be consistent with aqueous interpretations made previously for comparable landforms located at other candidate Martian crater-lake sites [e.g., Goldspiel and Squyres, 1991; De Hon, 1992; Forsythe and Blackwelder, 1998; Cabrol et al., 1996, 1998, 1999, 2001; Cabrol and Grin, 1999, 2001; Ori et al., 2000; Moore and Wilhelms, 2001; Newsom et al., 2003; Irwin et al., 2004; Bhattacharya et al., 2005; Di Achille et al., 2006; Mangold and Ansan, 2006]. Several viable aqueous mechanisms by which channel and valley features on Mars may have formed have been proposed in past studies, ranging from surface runoff under relatively warm and wet conditions to sapping processes under relatively cool and dry conditions [e.g., Milton, 1973; Masursky et al., 1977; Baker, 1982; Gulick and Baker, 1990; De Hon, 1992; Goldspiel et al., 1993; Craddock and Maxwell, 1993; Carr, 1995; Forsythe and Blackwelder, 1998; Carr and Malin, 2000; Gulick, 2001; Cabrol and Grin, 2001; Craddock and Howard, 2002; Aharonson et al., 2002; Grant and Parker, 2002; Hynek and Phillips, 2003; Carr and Head, 2003]. Although uncertainties remain regarding even the most basic Martian channel-formation processes, aqueous mechanisms are widely recognized as having the capacity to explain the nature of a wide spectrum of channel types on Mars, including all of the channels associated with the candidate crater-lake site examined here. Recently-discovered channel and fan systems with characteristics strikingly consistent with formation by aqueous processes [e.g., Malin and Edgett, 2003; Moore et al., 2003; Bhattacharya et al., 2005; Di Achille et al., 2006] may provide further support for fluvial and lacustrine interpretations previously made for the features of at least a proportion of candidate crater-lake sites on Mars (though it should be noted that such channel and fan systems also possess attributes consistent with development through non-aqueous processes).

3.2 Alternative Volcanic Interpretations of the Palos Crater / Tinto Vallis System

As noted for several other locales on Mars [e.g., Leverington and Maxwell, 2004; Greeley et al., 2005], the characteristics of some landforms of the candidate crater-lake site, including crater infill, the outlet breach, and the main inlet channel, are consistent with formation or substantial modification by igneous processes.

Infill materials at Palos crater appear to be layered and, in places, have lobate margins. The interior deposits of the central impact crater of the candidate lake site are essentially contiguous, by way of the outlet breach in the northwestern part of the crater rim, with smooth and ridged plains located southeast of Isidis Planitia. The wrinkle-ridged nature of these plains is consistent with that expected of volcanic flows [e.g., Hiller, 1979]. Wrinkle ridges are believed to form through horizontal shortening of near-surface layered deposits [Watters, 1988, 1991; Golombek et al., 1991; McGill, 1993; Schultz, 2000], and although the occurrence of wrinkle ridges on the terrestrial planets is not in principle restricted to volcanic materials [e.g., Schultz, 2000], all clear examples of these features are found in materials known or interpreted to be layered volcanic flows [e.g., Watters, 1988]. On the basis of the strong association between wrinkle ridges and volcanic plains on bodies such as the Moon and Venus [e.g., Young et al., 1973; Bryan, 1973; Muehlberger, 1974; Schaber et al., 1976; Schultz, 1976; Greeley, 1976; Whitford-Stark, 1981; Wilhelms, 1987; McGill, 1993; Basilevsky and Head, 1996], wrinkle ridges are used on Mars as features diagnostic of units likely comprised of layered volcanic flows [e.g., Potter, 1976; Greeley et al., 1977; King, 1978; Lucchitta, 1978; Peterson, 1978; Theilig and Greeley, 1979; Scott and Tanaka, 1986; Tanaka, 1986; Greeley and Guest, 1987; De Hon, 1992; Anderson et al., 2001; Head et al., 2002; Ivanov and Head, 2003; Hiesinger and Head, 2004].

The north breach in the rim of the main crater acted as a conduit for northward fluid flow, with liquids passing through the breach via a channel that cuts across plains that are contiguous with ridged plains that extend northwest toward Isidis Planitia [Hiller, 1979]. The channel is almost entirely filled with materials whose surface appearance is similar to that of adjacent ridged and non-ridged plains. If these materials are indeed volcanic, they could conceivably have accumulated in relation to flooding of the channel by volcanic flows derived from sources to the north, although the broad regional slope of Amenthes Planum is also consistent with flooding of the outlet channel by volcanic materials that flowed from the Tyrrhena Patera region toward Isidis Planitia.

The inlet system to the main crater consists in part of a complex dendritic network that is consistent with aqueous origins [e.g., see Craddock and Maxwell, 1993; Craddock and Howard, 2002] and is difficult to reconcile with igneous origins. Although very dense and complex channel networks can be formed through volcanic processes [e.g., Schaber, 1973; Leverington, 2004], localized highly dendritic systems that reach topographic divides lacking obvious volcanic sources are not strongly suggestive of formation by volcanic processes. However, the main inlet channel (Tinto Vallis) is comprised of a simple sinuous channel and tributary network that in the south are discontinuously expressed at the surface. The general morphology of the main inlet channel is notably consistent with that of volcanic rilles on the Earth, Moon, Venus, and Mars. These similarities are consistent on the basis of both overall appearance [Greeley, 1971ab; Howard et al., 1972; Young et al., 1973; Guest and Murray, 1976; Zisk et al., 1977; Wilhelms, 1987; Saunders and Pettengill, 1991; Baker et al., 1992, 1997; Head et al., 1992; Komatsu et al., 1992, 1993; Gregg and Greeley, 1993; Komatsu and Baker, 1994; Peterson et al., 1994; Kauahikaua et al., 1998; Leverington and Maxwell, 2004] and the general form and dimensions of high-resolution cross-sectional profiles [Wu et al., 1972; Swann et al., 1972; Strain and El-Baz, 1977]. These profiles, particularly with regard to systems such as lunar Rima Hadley, are broadly comparable to those described above in terms of appearance and width-to-depth ratios. The relatively shallow gradient (<1%) of the main reach of the sinuous inlet channel is also consistent with known lunar volcanic channels [e.g., Strain and El-Baz, 1977].

Importantly, it is not uncommon for volcanic rilles to be discontinously expressed at the surface as a result of incomplete collapse of channel roofs that, on the basis of terrestrial analogs [e.g., Peterson et al., 1994; Kauahikaua et al., 1998], are believed to form during eruptive events by processes that involve accretion on levees or aggregation of floating crustal rafts [Greeley, 1971a; Howard et al., 1972; see also, e.g., Keszthelyi, 1995]. Roofed sinuous volcanic channels that are defined in part by collapse pits are commonly found in association with large Martian volcanic rises, and examples of collapsed lava tubes that appear to have fed large volcanic flows are found on the flanks of Tyrrhena Patera [Greeley and Crown, 1990]. Beyond similarities related to the presence of collapse pits, it is notable that the relatively small volcanic rilles of the inner solar system commonly have simple channel forms that follow topographic slopes, have sharp rims, and have widths of up to several kilometers. Lunar, Venusian, and Martian volcanic channels are characterized by a wide range of sinuousities, and have lengths that can reach or exceed hundreds of kilometers [e.g., Baker et al., 1992; Komatsu et al., 1992, 1993; Komatsu and Baker, 1994; note that an alternative aqueous interpretation of the longest Venusian channels, the canali, has been suggested by Jones and Pickering, 2003]. Numerous volcanic channels have features often considered on the Earth to be characteristic of fluvial systems, such as levees, meander cutoffs, channel terraces, and streamlined islands [e.g., Greeley, 1971a; Baker et al., 1992, 1997; Head et al., 1992; Leverington, 2004]. Rilles widely interpreted as volcanic on the Moon and Venus are known to form complex dendritic or anastamosing systems that have a markedly fluvial appearance [e.g., Schaber, 1973; Baker et al., 1992; Komatsu et al., 1993; Leverington, 2004]. Examples of lunar rilles that cut across uplands are given below; the origins of rilles such as these are very poorly understood, but the geological settings of such features are dominantly volcanic, and the complete absence of hydrous minerals in lunar samples examined on Earth [Schmitt et al., 1970; Papike et al., 1991] strongly suggests that sinuous lunar rilles did not form in the presence of water.

The general geological context of the candidate crater-lake site is consistent with volcanic origins for certain site landforms. Specifically, the site is located immediately north of Tyrrhena Patera [McCauley et al., 1972; King, 1978; Plescia and Saunders, 1979; Greeley and Crown, 1990], one of the largest volcanoes on Mars. The flank materials located between the caldera of the volcano and the candidate crater-lake site consist of extensive ridged plains [e.g., King, 1978; Greeley and Guest, 1987; Greeley and Crown, 1990] that, on the basis of lunar analogs, are likely comprised of layered volcanic lava and ash deposits. The main inlet channel system of the site is nested within the fill of a northward-trending depression in local upland topography that connects the flanks of Tyrrhena Patera with the volcanic plains to the north. A channel origin strictly related to sapping or other aqueous mechanisms of formation is viable, but it also appears conceivable that the inlet channel formed partly as a constructional channel within volcanic flows erupted from Tyrrhena Patera. The interior deposits of most large impact craters in the region of the candidate crater-lake site are ridged plains, suggesting that local sources of volcanic flows may have also existed in the past, or that extensive volcanic flooding and subsidence have previously taken place.

Although theoretical and field-based studies have indicated that substantial amounts of incision can be associated with the flow of molten rock [e.g., Hulme, 1973; Peterson and Swanson, 1974; Cutts et al., 1978; Barnes and Barnes, 1990; Greeley et al., 1998; Kauahikaua et al., 1998, 2002, 2003; Wilson and Mouginis-Mark, 2001; Williams et al., 2004, 2005; see also Kerr, 2001], presumably through thermal and/or mechanical erosion [e.g., Fagents and Greeley, 2001], the actual mechanisms that might allow for incision across crater rims can only be speculated upon at present. Perhaps the simplest scenario would involve the exploitation by lava flows of existing zones of weakness in crater rims; such a mechanism would be similar to that hypothesized by, e.g., M.J. Grolier [in Masursky et al., 1978] to account for the incision of Rima Beethoven (Rima Prinz II) across a large upland barrier [see also Carr, 1974]. Regardless of the exact mechanisms involved, there are numerous examples of upland incision by lunar rilles that are both perplexing and directly relevant to the consideration of hypothetical volcanic mechanisms for incision into Martian crater rims. Lunar rilles that otherwise have the appearance of constructive mare lava conduits cut across local uplands in a manner that is not easily reconciled with incision at the relatively low present elevations of volcanic plains (e.g., Rima Beethoven, described in, e.g., Carr, 1974; Schultz, 1976; Strain and El-Baz, 1977; Wilhelms, 1987). Such rilles are hypothesized by some workers to have formed by the flow of lava at the terminations of volcanic highstands, with gradual incision taking place at the locations of lava conduits that were active during the lowering of magma levels, related to devolitization and phase change [e.g., Schultz, 1976; Greeley and Spudis, 1978].

It is conceivable that the upland terrain located at the foot of Tyrrhena Patera was flooded by lava flows in such a way as to form constructive channels in topographic lows and erosive channels at topographic divides. While mechanisms of formation are not understood, there are lunar examples of channels that appear to cut across crater rims at, e.g., Rimae Maupertuis at Montes Jura, and at what may be the rims of volcanic craters, including craters Krieger [Dietrich and Clanton, 1972; Greeley and Schultz, 1977; Wilhelms, 1987; Leverington and Maxwell, 2004] and Bowditch [El-Baz, 1972; West, 1972; Young, 1976; Leverington and Maxwell, 2004]. Such examples suggest that volcanic channels can, under appropriate eruptive conditions, cut across topographic divides in a manner analogous to that seen at Palos crater.

Figure 27: Hadley Rille (Rima Hadley), located on the lunar volcanic plains of Palus Putredinus, southeast Mare Imbrium [Apollo 15 Panoramic frame 9924; see also, e.g., Hackman, 1966] [Leverington, 2006]. Image center is at ~2.5oE, 25oN. Illumination is from the right.

Figure 28: Cross-sectional profiles of three lunar rilles: Rima Handel [after Strain and El-Baz, 1977]; Rima Beethoven [after Strain and El-Baz, 1977]; and Rima Hadley (Hadley Rille) [after Wu et al., 1972] [Leverington, 2006]. Profiles contain no vertical exaggeration. The width-to-depth ratios of lunar rilles can vary considerably within and between individual systems, with ratios of ~4:1 to 11:1 roughly describing the most common ranges that have been measured in past lunar studies based on high-resolution photogrammetric analysis of Apollo Panoramic-Camera images.

Figure 29: Oblique view of part of Rima Hadley (Hadley Rille), showing incomplete collapse of the roof of the channel [Apollo 15 Hasselblad frame AS15-87-11720; see also, e.g., Hackman, 1966] [Leverington, 2006]. Scale is valid only in the top-left corner of the image. Image center is at ~3.5oE, 26oN. Illumination is from image bottom.

Figure 30: Small lunar rilles with incompletely collapsed roofs (arrows), Marius Hills region of Oceanus Procellarum (Lunar Orbiter frame V-214-M) [e.g., Greeley, 1971a] [Leverington, 2006]. Image center is at ~56.75oW, 13oN. Illumination is from image top.

Figure 31: Sinuous lunar rilles that head in uplands east of lunar crater Plato (Lunar Orbiter frame IV-122-H3) [Leverington, 2006]. Channels extend onto the volcanic plains of Mare Imbrium [e.g., M’Gonigle and Schleicher, 1972]. Image center is at ~2oW, 49oN. Illumination is from the right.

Figure 32: Highly sinuous lunar volcanic channels. a: Rima Diophantus, located north of crater Diophantus in southwestern Mare Imbrium [Leverington, 2006]. Image center is at ~35o5, 28.75oN (Apollo 15 Panchromatic frame 299); illumination is from the right. b: Sinuous rille located near the southern foot of Montes Agricola, north of Aristarchus Plateau in Oceanus Procellarum. Image center is at ~53.25oW, 29.25oN (Apollo 15 Panchromatic frame 349); illumination is from image bottom.

Figure 33: Oblique view of a lunar rille (part of the Rimae Herigonius system) that cross-cuts highland terrain located northeast of crater Gassendi [Greeley and Spudis, 1978] (Apollo 16 Hasselblad frame AS16-119-19170) [Leverington, 2006]. Scale is valid only in the bottom-right corner of the image. The rille crosscuts highlands at ~36.75oW, 15.5oS. Illumination is from the left.

Figure 34: Two candidate crater-lake sites (L1 and L2; crater-lake sites 20.4S187.9 and 18.8S185.1 of Cabrol and Grin, 2001; see also Forsythe and Blackwelder, 1998) located west of Ma’adim Vallis (Viking MDIM mosaic; topographic data after Smith et al., 2003) [Leverington, 2006]. Igneous landforms in the region include extensive volcanic plains (e.g., at P) and a volcanic rise with a peak that is ~2400 m above surrounding terrain (V; see also Stewart and Head, 2001); possible volcanic origins for channels in such regions have not been previously considered in Martian paleolake investigations.

3.3 Discussion

Features of the examined candidate crater-lake site, such as localized dendritic networks associated with the western section of the inlet system, are consistent with formation by aqueous processes as previously hypothesized for other regions of the Noachian highlands [e.g., Craddock and Maxwell, 1993; Craddock and Howard, 2002]; formation through other mechanisms such as mass wasting cannot be ruled out. However, key landforms of the site, including the main inlet channel that extends from the flanks of Tyrrhena Patera, the fill materials of the central crater, and the outlet channel and associated fill materials, may be more simply accounted for by processes involving the voluminous flow of lava across the region. The roofed nature of the main inlet channel is consistent with the characteristics expected of relatively small volcanic channels of the inner solar system [e.g., Howard et al., 1972], and its interpretation as a sinuous volcanic rille is consistent with its range of width-to-depth ratios, its low gradient, and the presence of other previously-noted collapsed lava tubes on the flanks of Tyrrhena Patera [Greeley and Crown, 1990]. The infill materials of the central crater and outlet channel are contiguous with the ridged plains of Amenthes Planum that extend from the central crater toward Isidis Planitia, and that appear to be comprised of volcanic deposits related to the flow of lava. The site is located at the foot of a large volcanic rise, Tyrrhena Patera, and is very likely to have acted as a fluid conduit between this rise and the extensive volcanic plains to the northwest.

The feasibility of the volcanic model for formation of channel and infill features of the candidate crater-lake site rests heavily upon volcanic interpretations of lunar sinuous rilles used as analogs above. Beyond this basic consideration, the volcanic model represents a hypothesis that appears to be a reasonable alternative to otherwise compelling aqueous mechanisms of formation for landforms at the candidate crater-lake site. However, at present, there is not sufficient evidence to allow the volcanic hypothesis to be conclusively favored over proposed aqueous mechanisms, particularly in light of the existence of an overprint of localized dendritic channel networks associated with the western portion of the inlet system.

The existence at several Martian candidate crater-lake sites [e.g., Leverington and Maxwell, 2004; Greeley et al., 2005] of channels and crater infill with attributes that are consistent with volcanic origins suggests that other candidate crater-lake sites should be investigated with regard to the possible role of volcanic processes in their origin and evolution. Palos crater is a candidate crater-lake site that is located at the foot of a large volcanic rise, but the Memnonia site considered by Leverington and Maxwell [2004] is associated with extensive wrinkle-ridged plains that are not characterized by major volcanic constructs. Some candidate crater-lake sites on Mars combine the characteristics of these two cases and are directly associated with both broad wrinkle-ridged plains and relatively small volcanic rises. Many candidate crater-lake sites are located in highland regions such as Margaritifer Terra, Terra Meridiani, and Arabia Terra [e.g., Cabrol and Grin, 1999] where the widespread occurrence of ridged plains is strongly suggestive of massive Hesperian events of low-viscosity flood volcanism [see also, e.g., Greeley and Spudis, 1981; Tanaka, 1986]. Based upon lunar and Venusian analogs, the existence of ancient igneous channel systems at these and other volcanically-resurfaced regions should be expected. The role of volcanism in the formation and evolution of channels and basin fill at candidate crater-lake sites should be further investigated in the future in order to more clearly distinguish the effects of hypothesized aqueous processes from those of extrusive igneous processes.

As with the Memnonia system described above, hypotheses for the origins of channels and basin units in the Palos crater region are testable. Aqueous development would generally be expected to have resulted in widespread aqueous alteration of involved units. Volcanic development could conceivably have operated under drier conditions that favored preservation of original mineralogies. Volcanic development would also be expected to have resulted in emplacement of lava flows in basins located along local channel systems, as well as at the terminal basin of the system.

4 References Cited and Further Reading

Aharonson, O., M. T. Zuber, D. H. Rothman, N. Schorghofer, and K. X. Whipple (2002), Drainage basins and channel incision on Mars, Proc. Nat. Acad. Sci., 99, 1780-1783.

Anderson, R. C., J. M. Dohm, M. P. Golombek, A. F. C. Haldemann, B. J. Franklin, K. L. Tanaka, J. Lias, B. Peer (2001), Primary centers and secondary concentrations of tectonic activity through time in the western hemisphere of Mars, J. Geophys. Res., 106(E9), 20,563-20,586. Baker, V. R. (1982), The Channels of Mars, University of Texas Press, Austin, TX.

Baker, V. R., G. Komatsu, T. J. Parker, V. C. Gulick, J. S. Kargel, and J. S. Lewis (1992), Channels and valleys on Venus: Preliminary analysis of Magellan data, J. Geophys. Res., 97 (E8), 13,421-13,444.

Baker, V. R., G. Komatsu, V. C. Gulick, and T. J. Parker (1997), Channels and valleys, in Venus II, edited by S.W. Bougher, D.M. Hunten, and R.J Phillips, pp. 757-793, Univ. of Ariz. Press, Tucson.

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