Keywords: Planetary geomorphology, Fluvial channels, Volcanic channels, Mars, Venus, Titan

Table 1.

2.1. Lunar sinuous rilles

Channels visible by telescopic study of Earth’s moon (Fig. 1) initially looked promising as candidates for water flowing on that planetary surface (Pickering, 1904;Firsoff, 1960). The excellent, high-resolution images returned in 1966–1967 from the Lunar Orbiter missions that were designed to prepare for manned landing, provided data for detailed comparisons between sinuous lunar channel ways and terrestrial rivers (Peale et al., 1968;Schubert et al., 1970). Moreover, before the return of rock samples from the Moon, the prevailing theories held that, like Earth itself, Earth’s moon had a primordial, water-rich hydrosphere (reviewed byGilvarry, 1960). These theories even led the famous Nobel laureate chemist,Harold Urey (1967) to conclude that, given the obvious fluvial-like morphology of the sinuous rilles (e.g., their sinuous plan form), special processes must have occurred on the airless Moon to allow water to flow. He even suggested that a large comet impact might have produced a temporary water-rich atmosphere.Lingenfelter et al. (1968) used theoretical modeling to show that the lunar rivers could have been ice-covered and thereby able to flow. Following the consequences of this fluvial hypothesis,Gilvarry (1968) concluded that the lunar maria, instead of being the result of immense outpourings of lava, were actually the surface expressions of sediments and sedimentary rocks.

Geochemical analyses of rock samples returned from the Apollo landing missions showed clearly that the lunar primordial hydrosphere model was wrong and that the maria were the surface expressions of immense outpourings of basaltic lava (Taylor, 1982). The Moon clearly has no water-related context in which to place the sinuous rilles. In the lead up to the Apollo missions, the equifinality problem was encountered because a variety of nonaqueous processes were also hypothesized to be capable of producing sinuous rille morphologies. The hypothesized genetic mechanisms included lava channeling and the collapse of lava tubes (Baldwin, 1963;Oberbeck et al., 1969), pyroclastic flows (Cameron, 1965), and various combinations of structure and subsidence (Quaide, 1965;Schumm, 1970). In 1971, the Hadley Rille was inspected in the field by the Apollo 15 astronauts, and their findings, combined with studies of the regional geology showed that the lava channel and lava tube hypothesis was most consistent with all the available data (Greeley, 1971;Howard et al., 1972). Thus, in this case, the various advances in resolution and measurement during the intensive lunar exploration program of the 1960s eventually resolved the equifinality issue of water-versus-lava as the agent for forming the lunar sinuous rilles.

In retrospect, important morphological differences exist between fluvial channels or valleys and lunar sinuous rilles. Unlike fluvial forms, the latter most commonly have decreasing width and depth along their flow paths (Fig. 1). In a recent study of more than 200 lunar sinuous rilles,Hurwitz et al. (2013a) found that their lengths range from 2 to 566 km, with amedian width of about 500 m and amedian depth of about 50 m. These authors attribute the pervasive downstream decreases in width and depth to turbulent lava flow that facilitates thermal erosion in the proximal portions of the rille and then transitions distally to laminar flow, leading to a progressive decline in thermal erosion efficiency in a down-channel direction.

Hurwitz et al. (2012) used an analytical model to estimate the volume of lava needed to erode the Rima Prinz sinuous rill at ~50 km³ for a very low viscosity lava and about 250 km³ for an intermediate viscosity lava. From months to an Earth year would be required to form this feature by lava eroding at up to ameter per day, and the distal accumulation of solidified lava would extend over an area of about 2500 km². Thus, what was a highly erosive fluid near its source transitions to a much less erosive fluid, eventually leaving a huge solidified accumulation of all the channel-forming fluid. This is obviously very different than what would occur with water flows, which might leave some volume of sediment related to the amount of material eroded from the channel, but with most of the eroding fluid eventually leaving the depositional area by evaporation, infiltration, or other processes.

2.2. Mercury

The MErcury Surface Space ENvironment, GEochemistry, and Ranging (MESSENGER) mission that reached Mercury in 2008 revealed a dozen or so flat-floored, shallow valleys associated with extensive volcanic plains at high latitudes in Mercury’s northern hemisphere (Head et al., 2011;Byrne et al., 2013). The feature shown inFig. 2 is about 20 km wide with a smooth, lightly cratered floor that contrasts with the adjacent rough and highly cratered terrain into which it is shallowly incised. It was probably formed by the thermal and mechanical erosive action of high-magnesian, mafic or ultramafic lavas (Hurwitz et al., 2013b). The irregular knobs in the center right portion of the image are remnants of a rim material of the 140-km-diameter Kofi impact basin that was eroded by lavas emanating from vents about 50 km to the northwest of the image.

2.3. Io

Lava channels were observed on the surface of Jupiter’s moon, Io, in the course of the 1996–2003 imaging phase of the Galileo space mission (Keszthelyi et al., 2001).Schenk and Williams (2004) documented a particularly large channel, Tawhaki Vallis, that extends for about 200 km and is up to 6 km wide, representing either an ultramafic or a sulfur lava flow. The channel is shallowly incised (about 50 m) into plains that are probably primarily composed of sulfur (Schenk and Williams, 2004). The whole surface of Io must be relatively young, as it lacks impact craters. The young surface age and the lava that generated the channel derive from widespread volcanism that resurfaces Io through the intense tidal interaction of this relatively small moon with the massive planet Jupiter (McEwen et al., 2004).

3.1. Simple channels

Simple channels (Baker et al., 1992a;Komatsu et al., 1993) generally consist of a single, sinuous main channel that lacks the complex branching and anastomosing reaches characteristic of other varieties of Venusian channels (Gulick et al., 1992a). Simple channels can be further subdivided into simple channels with flow margins, sinuous rilles, and canali (Fig. 3). Some simple channels are located on well-defined flow deposits or flow fields (Fig. 3A) (Komatsu et al., 1993). These simple channels with flow margins are similar in morphology to channels that form on terrestrial lava flows. Because these channels have formed on lava flows and do not seem to have incised surrounding terrain, they appear to be similar to their terrestrial counterparts in being mostly constructional in origin. In general, these simple channels lack distinctive source regions.

Sinuous rilles emanate from distinct, circular, or elongated regions of collapse (generally several kilometers in diameter), and they form channels up to several kilometers wide and tens to hundreds of kilometers long (Baker et al., 1992a;Komatsu et al., 1993). As in the case of lunar sinuous rilles, these Venus counterparts become narrower and shallower in a downstream direction (Fig. 3B). Most sinuous rilles on Venus are not associated with detectable lava flow margins. The similarities in morphology and size to lunar sinuous rilles may imply that thermomechanical erosion by high-discharge, highly fluid lava was also an important channel-forming process on Venus (Komatsu et al., 1993;Komatsu and Baker, 1994a;Oshigami et al., 2009). Some of the Venusian sinuous rilles are associated with networks of valleys or depressions (Baker et al., 1992a;Gulick et al., 1992a,b;Komatsu et al., 1993,2001;Oshigami et al., 2009).

Canali are features that are unique to Venus. Unlike other simple channels they have generally constant width and depth over their entire flow path (Fig. 3C) (Baker et al., 1992a,1997;Komatsu et al., 1992,1993). These channels generally have widths ranging up to 3 km and lengths exceeding 500 km. However, some canali may be up to 10 km wide; and a few have enormous lengths, up to 6800 km in the case of Baltis Vallis (Baker et al., 1992a;Komatsu et al., 1993). Canali may locally exhibit numerous abandoned channel segments, cutoff meander bends, levees, and radar dark terminal deposits (Baker et al., 1992a,1997;Komatsu et al., 1993;Kargel et al., 1994). Sources and termini are generally indistinct. Canali are generally located on topographic plains (Komatsu et al., 1993), considered to be volcanic in origin and mafic in composition (Kargel et al., 1993), and they are tectonically deformed along their longitudinal profiles (Komatsu and Baker, 1994b;Langdon et al., 1996). The extraordinary length and a relatively short formation time scale (i.e., geologically speaking) of Baltis Vallis allowed this feature to be used to correlate distant geological units on Venus plains to understand their sequential relationships (Basilevsky and Head, 1996).

The morphology of these channels suggests that they probably formed by continuously conveyed large discharges of low viscosity lava to distant regions over prolonged periods (Baker et al., 1992a,1997;Komatsu et al., 1992,1993;Bray et al., 2007). Both erosional and constructional origins have been proposed (Komatsu et al., 1992;Gregg and Greeley, 1993;Bussey et al., 1995;Williams-Jones et al., 1998;Lang and Hansen, 2006;Oshigami and Namiki, 2007). The formative fluids have been hypothesized to be a wide range of silicate lava varieties, as well as low-viscosity flows of sulfur or carbonatite lava (Baker et al., 1992a,1997;Komatsu et al., 1992,1993;Gregg and Greeley, 1993;Kargel et al., 1994). More speculative hypotheses invoke the role of non volcanic fluids, including turbidity currents that would have had to occur at a time when an ocean existed on the Venusian surface (Jones and Pickering, 2003). Alternatively,Waltham et al. (2008) envisioned particulate gravity currents resulting from the suspension of fine particulate matter in the dense Venusian atmosphere and moving downslope to travel long distances. The origin of the Venusian canali remains poorly understood.

3.2. Complex and compound channels

Complex channels form anastomosing, braided, or distributary patterns that are generally (but not always) on flow deposits (Gulick et al., 1992a;Komatsu et al., 1993). Individual channel widths range from ~3 km down to the limit of resolution, while the total width of the channel system can range from 20 to 30 km, and up to hundreds of kilometers in length (Gulick et al., 1992a). Most complex channels are located on flow deposits and are classified as complex channels with flow margins (Fig. 4A) (Komatsu et al., 1993). Complex channels are often located along with simple channels on flow deposits, indicating a genetic connection to the lava flows. These channels are commonly separated by radar-bright (or radar-dark in some cases) material that is considered to be lava, and the channels probably formed by a constructional process. Complex channels that are not located on flow deposits appear to have eroded into surrounding terrain. This particular subclass of complex channels is simply known as complex channels without flow margins (Fig. 4B) (Komatsu et al., 1993).

Compound channels contain simple and complex segments (Fig. 4C) (Komatsu et al., 1993). The channels vary greatly in size, with widths ranging between several tens of kilometers in complex regions down to the limit of resolution in simple reaches. Total lengths of compound channels can range from 75 km to thousands of kilometers (Gulick et al., 1992a).

Kallistos Vallis (Fig. 4C) is a particularly interesting compound channel. It emanates from a distinct collapse region (Baker et al., 1992a,1997;Komatsu et al., 1993). However, instead of becoming distally narrower and shallower like a sinuous rille, the channel displays a great variety of morphologies as it extends about 1200 km. Some morphologic characteristics of Kallistos Vallis, in particular the collapsed source region and anastomosing segment, bear a resemblance to some aspects of what have been termed Martian out flow channels, and the informal nameoutflow channel was given to Kallistos Vallis (Baker et al., 1992a).Leverington (2011) has drawn particular attention to these features of Kallistos Vallis and applied theoutflow designation much more generally to lunar and Venusian sinuous rills and various other volcanic channel forms sourced at fissures, vents, collapse areas, and other types of depressions, proposing further that these all share a common origin with the Martian channels (see Section 7).

Important differences can be documented between Kallistos Vallis and what are more properly termedcataclysmic flood channels on Mars. Of course, the regional geological context is totally different. Venus is a planet dominated by the manifestation of volcanic processes, with the greatest variety of lava-related flow features to have yet been discovered. No regional or temporal context indicates a role of water in forming any of the landscape features on the surface of Venus. Mars is a completely different planet in regard to the role of water, with rich manifestations of relict fluvial forms and other indicators of water and water–ice compositions and processes in shaping the planetary surface (see Sections 5 and 7). Ancient Mars was earthlike in that regard (Carr, 2012); Venus was not.

The main similarities of Kallistos Vallis to anoutflow channel are its collapsed source region (upper left ofFig. 4C) and the anastomosing segment (lower center ofFig. 4C). However, the collapse region leads not to a fluvial channel but to a linear trough about 400 km long and about 600 km deep. The collapse source pit is connected to the trough by an incised gorge, and a sinuous canali-type channel also emanates from this gorge. The canali channel is about 1.5 km wide and 175 km long. Typical for Venusian canali, it maintains a relatively constant width over its entire length, implying genesis under very poorly understood conditions by a poorly understood fluid process. The linear trough of Kallistos Vallis eventually narrows to only 1.5 km wide and shallows to less than 150 m deep. Beyond this point the channeled fluid seems to have spilled out to create the distinctive anastomosing sub channels that are spread over a width of as much as 18 km (lower center part ofFig. 4C). Deflected eastward, the flows were impounded upstream of a north–south ridge, eroding through that obstacle to create streamlined hills in the divide crossing. Downstream of this divide the system displays a distinctive distributary pattern of radar-bright channels (lower right ofFig. 4C) that feed into an immense area of lobate deposits, the likely solidified flows that traversed the channel, and these cover an area of about 100,000 km². The distributary pattern of channels and most of the lava plains are not shown inLeverington’s (2014) map of the system.

4.1. Fluvial drainage distribution

Drainages mapped in SAR data (Lorenz et al., 2008;Langhans et al., 2012;Burr et al., 2013a) show a near-global distribution, although their density is not homogeneous. A parameter entitleddelineated fluvial feature density was calculated for each band of 30° latitude as the total distance along delineated network links ratioed to the area of SAR coverage. The parameter is similar to drainage density but, because the resolution and noise of the SAR data preclude delineation of the low-order links or small networks, it underestimates true drainage density. Comparison of values among latitudinal bands shows that fluvial features are denser, by an order of magnitude, at high northern latitudes (Burr et al., 2013a) where they drain radar-bright, rugged terrain and empty into the numerous north polar lakes (Stofan et al., 2007;Hayes et al., 2008;Cartwright et al., 2011). Like the lakes, the north polar networks are commonly radar-dark, possibly as a result of a drape of fine-grained organic sediments deposited either during back flooding of the river valleys during lake high-stand or during the waning stage of fluvial flow (Lorenz et al., 2008;Burr et al., 2013b).

Fluvial features are also concentrated around the Xanadu province (Burr et al., 2013a), a photometrically and geographically distinct region that stretched over ~100° along the equator (Radebaugh et al., 2011). Standing as high as 2000 m above the surrounding landscape, Xanadu (like the north polar regions) has radar-bright high-relief terrain with irregular orcrenulated texture, inferred to be mountain ranges (Radebaugh et al., 2011). Radar-bright fluvial valleys with cobble-sized sediments are incised within the crenulated terrain (Fig. 7) (Burr et al., 2013b). The fact that most of these networks within Xanadu are below the resolution of the SAR data suggests that other uplands on Titan that lack discernable networks may nonetheless be fluvially dissected (Burr et al., 2013b). The networks imaged by DISR, although not discernable in overlying SAR data, are evidently heavily incised. Some higher-order links and trunk valleys are visible, as are wide radar-bright extensions that stretch from the Xanadu networks across the surrounding plains (Le Gall et al., 2010). The radar-bright return from these wide fluvial features is interpreted as a result of rounded cobbles greater than a few centimeters in diameter (Le Gall et al., 2010), and the features themselves are inferred to be the deposits of gravel-bed, braided, ephemeral rivers (Burr et al., 2013b).

The remaining fluvial networks visible in SAR data are scattered within the mid-latitudes (Burr et al., 2013a,b); the extensive tropical dunes on Titan apparently either preclude the formation of fluvial flow because of aridity and/or cover over past fluvial flow features at tropical latitudes (Lorenz et al., 2006;Radebaugh et al., 2008;Lorenz and Radebaugh, 2009). In contrast to the polar and Xanadu networks, these scattered mid-latitude networks occur in relatively low-relief settings (theundifferentiated plains ofLopes et al., 2010), where they commonly form broad, shallow, radar-dark features, hypothesized to be braided or anabranching channel patterns (Lorenz et al., 2008;Burr et al., 2013b). Often radar-dark, they are interpreted to be braided or anabranching fluvial systems with possible terminal splays of fine-grained sediment (Lorenz et al., 2008;Burr et al., 2013b).

4.2. Fluvial sediments observed on the surface

Although the networks imaged by DISR are below the resolution of the SAR data (Soderblom et al., 2007), their appearance in the DISR images provides some indication of sediment type and size. In these visible-wavelength images from altitude, the DISR networks appear dark and are interpreted as being mantled with fine-grained, likely organic, sediment (Tomasko et al., 2005;Perron et al., 2006). The Huygens probe landed several kilometers from the outlet of the fluvial network on a dark flat surface hypothesized to be a dry lakebed. At the surface, the DISR camera images show rounded cobbles (Fig. 8), inferred to be icy but with significant non-icy material. The relationship between the fluvial networks and the icy cobbles is not clear, but the presence of the cobbles provides plausibility for the interpretation of rounded icy cobbles in the radar-bright fluvial features draining Xanadu (Le Gall et al., 2010).

4.3. Drainage (or fluvial) network morphologies

By virtue of their regional extent and responsiveness to formative conditions, drainage or fluvial networks provide important evidence for surface, and subsurface conditions. Terrestrial drainages have been classified into several basic and modified patterns (e.g.,Howard, 1967; although seeDrummond, 2012, for a discussion of the slight differences among drainage classification schemes). Each of these drainage patterns carries specific implications regarding landscape slope, substrate erodibility biases, and other conditions at the time of fluvial runoff. Based on a quantitative algorithm for classifying drainages on Earth, an analysis of drainage morphologies visible in SAR and DISR images classified one-half of all the mapped networks as rectangular (Burr et al., 2013a). These rectangular networks are globally distributed in the available SAR data (Burr et al., 2013a), although gaps in the SAR coverage preclude a rigorous statistical analysis of geospatial distribution. As rectangular networks on Earth are commonly the result of control on overland flow by subsurface structures or topography associated with structures, this finding for the Titan networks was interpreted as indicating wide-spread structural control on fluvial flow (Burr et al., 2013a).

5.1. Timing of fluvial activity

The extensive dissection of the heavily cratered terrain on Mars by valley networks (e.g.,Fig. 10) was long used to argue that the networks themselves dated to the Noachian epoch (e.g.,Carr and Clow, 1981;Carr, 1996). As with other age categories for Mars, the Noachian epoch is defined by the density of impact craters and by comparisons to radiometric dates on lunar cratering (e.g.,Hartmann and Neukum, 2001). This procedure defines the Noachian as the portion of Mars geological history prior to 3.7 Ga. Later Mars epochs are then divided into the Hesperian, from about 3.7 to 3.0 Ga, and the Amazonian for surfaces younger than 3.0 Ga.

Recent work has shown that the formation of the well-developed valley networks on Mars is more concentrated in time, with much of the activity occurring close to the Noachian/Hesperian boundary (Howard et al., 2005;Irwin et al., 2005b). Moreover, as was apparent from some of the older Viking imagery of Mars (Baker and Partridge, 1986), fluvial activity in the valley networks continued into the Hesperian (Mangold and Ansan, 2006;Bouley et al., 2009,2010;Hynek et al., 2010). Also apparent from the earlier Viking images, dense networks of fluvial valleys dissect Amazonian-aged surfaces on some Martian volcanoes, such as Alba Patera and Hecates Tholis (Gulick and Baker, 1989,1990).

5.2. Valley networks

As noted inTable 2, two major types of Martian valleys can be distinguished. The longitudinal valleys (Baker, 1982) are relatively wide and elongate with few tributaries. They commonly dissect upland plateaus, and their theater-like valley heads suggest an important role of groundwater seepage undermining slopes (i.e., sapping, in their origin;Goldspiel and Squyres, 2000;Harrison and Grimm, 2008). Examples include Nirgal Vallis and Nanedi Vallis. Some of these valleys have small but relatively well-preserved deltas at their termini (Fassett and Head, 2005;Irwin et al., 2005a;Mangold and Ansan, 2006;Di Achille et al., 2007;Mangold et al., 2007;Kraal et al., 2008;Hauber et al., 2009;Dehouck et al., 2010).

Topographic data provided by the Mars Orbiter Laser Altimeter (MOLA) instrument on the Mars Global Surveyor (MGS) orbiter (Smith et al., 1999) show that the orientations of the numerous multi-branched networks of valleys are consistent with gravitational control of fluid flow on the Martian surface. The latter is locally warped by the formation of a trough and bulge that formed around the immense loading of the crust by the Tharsis volcanics (locality TV,Fig. 9) in late Noachian time (Phillips et al., 2001).

The network valleys widen and deepen in the downstream direction (Craddock and Howard, 2002;Howard et al., 2005;Irwin et al., 2008;Hynek et al., 2010). Small channels that are relicts from the rivers that formed the valleys are only rarely discernable because of the relatively young eolian deposits that commonly mantle the valley floors (Irwin et al., 2005a;Jaumann et al., 2005;Kleinhans, 2005). Similarly, deposits at the valley termini, are commonly missing, either because of erosion or mantling by younger lava flows, mostly of Hesperian age (Irwin et al., 2005a;Ansan et al., 2008;Ansan and Mangold, 2013).

The image resolution issue, discussed in Section 1, played an important role for valley network interpretation in that the relatively low-resolution images from the Mariner 9 and Viking Orbiter missions of the 1970s and 1980s seemed to indicate low drainage densities (length of valleys or channels per unit area) for the networks. Based on a global analysis of the relatively low-resolution Viking orbital imageryCarr (1996) andCarr and Chuang (1997) inferred that areas of highly dissected southern highlands on Mars had average drainage densities of only ~0.005 km⁻¹, which is much lower than typical values for fluvial drainage on Earth. These low values suggested that regional rainfall and runoff processes might not be the major cause for valley network formation on Mars (e.g.,Squyres and Kasting, 1994;Segura et al., 2002).

The early summaries of drainage densities on Mars did not incorporate some results, also from Viking data (Baker and Partridge, 1986), that showed that local valley networks in the heavily cratered terrain consist of younger (pristine) elements that are portions of much older, though degraded networks. By considering the latter,Baker and Partridge (1986) found that the degraded network densities were as high as 0.1 km⁻¹, which is consistent with some terrestrial data. Using higher-resolution Viking data,Gulick and Baker (1989,1990) determined that drainage densities were actually an order of magnitude higher for the fluvial valleys that formed on some younger Martian volcanoes (most notably Alba and Hecates) than those in the heavily cratered Noachian terrains. These values were much more similar to their terrestrial counterparts, with the Alba valleys (locality B,Fig. 9) having values between 0.3 and 1.5 km⁻¹ compared to 0.2 and 5.0 km⁻¹ on the Hawaiian volcanoes (Gulick and Baker, 1989,1990). These higher values imply at least localized atmospheric precipitation (Gulick et al., 1997), and overland flow developed on a volcanic ash mantle overlying the very porous volcanic lava flows (Gulick and Baker, 1990).

The subsequent recognition of higher drainage densities using MOLA topographic data and higher resolution imagery from the MGS mission of the late 1990s (e.g.,Hynek and Phillips, 2003) provided a confirmation of what had been shown in the more localized study of Viking data, and it is now clear that drainage densities average 0.1 to 0.2 km⁻¹ over extensive areas of the Martian surface (Ansan and Mangold, 2006,2013;Luo and Stepinski, 2009;Hynek et al., 2010). The high values of drainage density strengthen the case that prolonged precipitation and runoff processes were necessary for the origin of the valley networks (Ansan and Mangold, 2006;Craddock and Howard, 2002;Mangold et al., 2004;Quantin et al., 2005;Mangold and Ansan, 2006;Ansan et al., 2008;Hynek et al., 2010).

Mangold et al. (2012) relate the evolution of fluvial landscapes in the heavily cratered terrains of Mars to degradation of the highland craters. Craters older than about 3.9 Ga (Middle Noachian) date from the Late Heavy Bombardment, a pulse of very high impact fluxes that occurred throughout the inner solar system for about 100 My around 3.9 Ga. These ancient craters typically have heavily degraded rims, and the older ones are essentially rimless, with their ejecta having been completely eroded away, probably by fluvial processes. The eroded materials fill many of the crater floors and may represent a period of prolonged and effective fluvial planation (Howard et al., 2005;Irwin et al., 2005b). In contrast, the valley networks developed near the time of the Noachian/Hesperian transition (~3.7 Ga) on this planation surface. They dissect areas around craters that are eroded, but much less so than the more ancient rimless forms. In the Hesperian, from about 3.7 to 3.3 Ga, further degradation occurred within craters, leading to alluvial fans on their floors (Moore and Howard, 2005) and local dissection of rims but general preservation of the ejecta. Amazonian craters (younger than ~3.3 Ga) lack fluvial landforms and are relatively fresh in appearance with pristine-looking ejecta blankets and central peaks.

Detailed work in the Libya Montes area, just to the southeast of the Syrtis Major volcanic complex (locality Y,Fig. 9), shows the later evolution of Martian valley networks in relation to standing bodies of water on the planet’s surface. The dendritic valley networks in the region were formed in the Noachian between about 4.1 and 3.8 Ga (Jaumann et al., 2010;Erkeling et al., 2012), with some activity continuing into the Hesperian. The fluvial activity was associated with the ponding of water in craters with associated deltas, hydrated minerals, and alluvial fans. At the western end of the Libya Montes, near its margin with the Syrtis Major volcanic complex, a large valley system, Zarga Vallis, has (i) an older, eastern network of dendritic valleys that probably formed by precipitation runoff processes and (ii) a younger, western segment that is a longitudinal valley that probably developed by volcanic melting of ground ice (Jaumann et al., 2010). The transition from runoff valley development to sapping or ground-ice melting seems to have occurred in the middle Hesperian (~3.6 Ga).

5.3. Alluvial rivers, deltas, and sedimentary rocks

On 6 August 2012, the Curiosity Rover of the Mars Science Laboratory Mission successfully landed on the floor of the 150-km-diameter Gale Crater (locality G,Fig. 9). The material on which it landed was a fluvial conglomerate (Fig. 11), deposited near the distal end of an alluvial fan (Fig. 12). As they were reported in the popular media, these features had the appearance of unique discoveries. However, the landing site had actually been carefully chosen to make such observations; the choice was based on the developing understanding of Mars’ watery early history. Combined with observations from the Mars Exploration Rover (MER) Opportunity landing site (theBurns Formation ofGrotzinger et al., 2005,2006) and the related documentation of sedimentary rocks by various orbiters, the recent lander studies leave no doubt that Mars had a watery ancient past, involving the extensive emplacement of sedimentary rocks (Grotzinger and Milliken, 2012).

Until the later 1990s, channels and valley networks cut into rock had comprised the main evidence that was cited in support of the view that Mars once had conditions that supported an earthlike hydrological cycle. The view of a water-rich planet sharply contrasted with current conditions on the planet and with the then-prevailing views from physics and chemistry that Mars had always been cold and dry (Baker, 2014c). However, starting in the late 1990s, a rapid succession of discoveries added to the list of evidence for more earthlike hydrological conditions on early Mars. Although possible delta and fan-like depositional landforms had been tentatively recognized from the older low-resolution data (e.g.,De Hon, 1992;Cabrol and Grin, 1999;Ori et al., 2000), the meter-scale images of the Mars Obiter Camera (MOC) on the MGS spacecraft led to the key discovery of deltaic features very similar in general morphology to what occurs on Earth (Malin and Edgett, 2003). The most impressive of these discoveries is a delta (Fig. 13) in Eberswalde Crater (locality U,Fig. 9). Multiple studies have estimated a dominant, channel-forming discharge for the paleochannels on this delta at a few to several hundred cubic meters per second (Malin and Edgett, 2003;Moore et al., 2003;Jerolmack et al., 2004;Howard et al., 2007;Irwin et al., 2014).

Related to the Eberswalde Delta discovery (Malin and Edgett, 2003) is the recognition of sinuous channels showing meander cutoffs, scroll topography, and related features of alluvial rivers with floodplains. Unlike rivers that are cut into bedrock, alluvial rivers on Earth have channel beds and boundaries composed of the same sediments that they transport. The Mars alluvial channels commonly display an inverted relief, probably because eolian deflation selectively removed the fine-grained sediments of the adjacent floodplains relative to the coarse-grained channel-filling sediments (Williams and Edgett, 2005;Pain et al., 2007). Particularly extensive alluvial river paleomeander belts occur in the Aeolis Dorsa region (Burr et al., 2009a,2010;Williams et al., 2009a,2013), which lies a few hundred kilometers east of the Gale Crater region (locality G,Fig. 9). The meandering patterns seem to have formed in alluvial rivers because of the deposition of relatively fine sediments (clay/silt muds) that were flocculated by dissolved salts (Matsubara et al., 2014).

Another interesting Mars delta (Fig. 14) occurs in Jezero Crater in the Nili Fossae region of Mars (locality F,Fig. 9). Hyperspectral data from the Visible and Near Infrared Mineralogical Mapping Spectrometer (OMEGA—Observatoire pour la Minéralogie, l’Eau, les Glaces et l’Activité) instrument on Mars Express and the Compact Reconnaissance Imaging Spectrometer (CRISM) instrument on the Mars Reconnaissance Orbiter (MRO) revealed that this region has a diverse assemblage of minerals, including phyllosilicates (e.g., clay minerals), consistent with widespread liquid water activity that range from surface weathering to hydrothermal processes (Mangold et al., 2007). A valley network feeds from this altered source terrain to bedded sediments on the floor of Jezero Crater (Fig. 14), and these were probably emplaced in a crater lake near the time of the Noachian/Hesperian boundary (Fassett and Head, 2005). Some of these sediments are rich in iron–magnesium smectite clay (indicated in green on the image shown inFig. 14) (Ehlmann et al., 2008), and the clay was probably transported as suspended load by a river that drained source areas characterized by clay-rich rocks (Mangold et al., 2007;Ehlmann et al., 2008).

While most Mars delta landforms are younger than the main phase of valley network development, a striking example of a late Noachian/early Hesperian delta occurs at Terby Crater, a 174-km-diameter feature on the northern edge of Hellas Planitia (locality H,Fig. 9) centered at 28°S, 73°E. At that location a cumulative thickness of 2000 m of sedimentation is inferred (Ansan et al., 2011). Such great accumulations of sediment show that the erosion by the valley networks probably was an important contributor to the filling of craters in the ancient heavily cratered terrain of Mars.

5.4. The early Mars climate conundrum and a northern plains ocean

The spatial distribution of the valley networks is not uniform throughout the heavily cratered terrain (Gulick, 2001), as might be expected if impacts into the southern highlands were responsible for episodic formation of the networks (e.g.,Toon et al., 2010). Instead, the densest concentrations of valleys (Fig. 9) follow a swath that circles the planet, extending several hundred kilometers into cratered highlands from the latter’s boundary with the northern plains. Allowance must be made for the great Tharsis volcanic province, which is younger than the valley networks and which imposed itself on this boundary. This pattern is consistent with what would be expected for a precipitation source associated with a hypothetical northern plains ocean (Luo and Stepinski, 2009;Stepinski and Luo, 2010). The relationship of valleys to the northern margins of the heavily cratered terrain of the southern highlands has been confirmed by an independent study that quantified the spatial distribution of drainage densities (Hynek et al., 2010). The distribution also corresponds to the locations of deltas that are graded to the base level of the northern plains ocean (Di Achille and Hynek, 2010), to the concentrations of sedimentary rocks on Mars (Malin and Edgett, 2000a;Delano and Hynek, 2011), and to the presence of high-Al clay minerals in deep weathering profiles (Le Deit et al., 2012;Loizeau et al., 2012). The latter would require intense leaching in a surficial environment for their formation. Moreover, such a leaching episode would occur at the Hesperian/Noachian boundary (Loizeau et al., 2012), corresponding to the same episode of precipitation that is recognized in regard to the valley networks.

An ancient ocean-scale water body, about 3 × 10⁷ km² in area, informally namedOceanus Borealis (Baker et al., 1991), has long been hypothesized for the northern plains of Mars (Fig. 9). Although it was initially inferred from the mapping of sedimentary landforms (Jons, 1985;Lucchitta et al., 1986), the ancient ocean hypothesis was more controversially tied to the identification of what were interpreted as shoreline landforms byParker et al. (1989,1993). However, failure to confirm the latter (Malin and Edgett, 1999,2001) and variations in the hypothesized shoreline elevations of up to a couple of kilometers (Carr and Head, 2003) led some to reject the hypothesis. More recent studies, including the explanation of shoreline deformation by true polar wander generated by the formation of the immense Tharsis rise (Perron et al., 2007) and interpretations of compositional data (Dohm et al., 2009), have lent support to the ocean hypothesis. Moreover, recent studies have shown that, unlike Earth’s oceans, Mars’Oceanus Borealis formed episodically (Baker et al., 1991;Fairén et al., 2003). At least one earlier oceanic phase occurred at a time coincident with the formation valley networks in the Martian highlands (Clifford and Parker, 2001), and later phases were associated with inflows from the immense cataclysmic flood channels that are described in Section 7.

Multiple studies indicate that the formation of the valley networks required prolonged periods of rainfall in amounts comparable to what occurs for Earth’s arid or semiarid regions (Howard, 2007;Barnhart et al., 2009;Luo and Stepinski, 2009;Hoke et al., 2011;Irwin et al., 2011;Matsubara et al., 2013). Associated lakes, deltas, and alluvial fans show complex histories of fluctuating water and sediment discharges (Malin and Edgett, 2003;Moore and Howard, 2005;Fassett and Head, 2005,2008;Di Achille et al., 2006;Pondrelli et al., 2008;Di Achille and Hynek, 2010;Grant et al., 2011;Hoke et al., 2014), and these also imply prolonged periods of precipitation and runoff (Moore et al., 2003;Jerolmack et al., 2004;Matsubara et al., 2011).

This extensive evidence for warm, wet conditions on early Mars climate poses a conundrum because of the inability of current theoretical models to explain all this. In contrast to earlier theory (e.g.,Pollack et al., 1987), if one assumes a CO₂–H₂O atmosphere, which involves the most parsimonious extrapolation that can be made from current Mars conditions to those of the ancient past, then brining the global mean surface temperature of Mars to near the freezing point of water would not physically have been possible at or before about 3.8 Ga ago (Kasting, 1991). These same assumptions apply to more complex calculations (Forget et al., 2013;Wordsworth et al., 2013), which also lead to a conclusion that a CO₂–H₂O Martian atmosphere cannot generate the needed warm temperatures.

Another view holds that it may not be necessary to bring global mean temperatures above the freezing point.Gulick et al. (1997) explored the potential climatic effects of instantaneous pulses of CO₂, such as may have been released during outflow channel formation and subsequentocean formation as hypothesized byBaker et al. (1991). They found that a one to two bar pulse is sufficient to raise mean global temperatures above 240 or 250 K for tens to hundreds of millions of years, even when accounting for CO₂ condensation. Such pulses can place the atmosphere into a stable, higher pressure, warmer greenhouse state, where substantial water volumes could be transported from a frozen lake or ocean to higher elevations, despite global temperatures well below freezing. This water, precipitated as snow, could melt, infiltrate, and runoff, ultimately forming fluvial valleys in the southern highlands if associated with localized heat sources and hydrothermal systems, such as magmatic intrusions, volcanoes, or cooling impacts (Gulick, 1998). Thus, if outflow channel discharges were accompanied by a significant release of CO₂, a limited hydrological cycle could result that would be capable of producing fluvial erosion and valley formation. Sources of atmospheric CO₂ during this time could have been provided by venting from associated volcanism, release of gases dissolved in groundwater, de-adsorption of gases from inundated regolith, and vaporization of clathrate in the regolith and ices resident in the polar caps (Baker et al., 1991).

Mischna et al. (2013) recently proposed another scenario. They envision combinations of three driving factors for promoting transient warm/wet conditions on early Mars: (i) an insolation effect, mainly driven by changes in Mars’ obliquity; (ii) a trigger effect, mainly as it will subsequently promote a transient water-rich greenhouse effect; and (iii) an albedo effect involving relatively dark portions of the Martian surface. The insolation effect results in periods of increased solar heating at various latitudes. The albedo effect can arise either (i) from dark, dust-free exposures of basalt bedrock or (ii) from the temporary presence of relatively low albedo, ponded water, notably the northern plains ocean inferred for early Mars, acting in the manner hypothesized byBaker et al. (1991) andBaker (2001,2009a). Finally, a trigger effect can be provided by the massive, short-termvolcanic injection into the atmosphere of particularly potent greenhouse gases, such as sulfur dioxide. Though such gases may be generally short-lived in the atmosphere, their temporary warming effect can provide a trigger to get large quantities of water vapor into the atmosphere and that water will contribute to more prolonged greenhouse warming.

Halevy and Head (2014) also invoke the episodic volcanic release of sulfur dioxide, but in combination with aerosols, as a means of short-term warming of a dusty Martian atmosphere. The combination of mechanisms envisioned byMischna et al. (2013) might then achieve the necessary warming, especially adjacent to the margins of the northern plainsocean that would supply water vapor. The warming could persist long enough to generate the rainfall/runoff conditions needed to produce the valley networks, but ultimately Mars would return toice-house conditions as the obliquity changed on timescales of millennia. These overall interactions are similar to what was hypothesized byBaker (2009a).

5.5. Gullies and other recent flow phenomena

The discovery of gully forms (Malin and Edgett, 2000b) in MGSMOC images sparked a spirited new debate over the history of water on Mars. Although current average temperatures are below 273 K and atmospheric pressures are at or below the triple-point vapor pressure of water at 6.1 mbar, many investigators concluded that the morphology of at least some of the gullies implied a formation mechanism involving the flow of water. Additional studies imply that some gully modification processes may still be ongoing today, including the discovery of gullies on surfaces devoid of craters and gullies with deposits that overlap other modern, possibly active, landforms, such as dunes and polygons.

Although terrestrial gullies are commonly defined simply by the presence of an incised channel segment,Malin and Edgett (2000b) defined the Martiangullies as having an alcove in the source region, a defined channel or system of channels in the mid-section and a debris apron in the terminus. The formation and modification of gully systems on Earth involve a variety of fluvial and hillslope processes. These include fluvial (including overland, soil water, and groundwater flow), colluvial, and mass wasting (e.g., debris flows and avalanches) processes. The challenge posed by the Martian gullies is to determine which of these, or other, mechanisms are primarily responsible for gully formation and which are simply modification processes.

6.1. Channel structures within volcanic plains

On Mars, basaltic provinces divide into three main types (Greeley and Spudis, 1981;Zuber and Mouginis-Mark, 1992): (i) large shield volcanoes fed primarily from centralized sources (e.g., Olympus Mons and the other Tharsis Montes;Bleacher et al., 2007), (ii) regional packages of sheet-likeflood lava that appear to originate from fissures (e.g., the Cerberus Fossae Units in Elysium Planitia;Plescia, 1990;Jaeger et al., 2007,2010), and (iii) vent fields composed of coalesced lava from widely distributed low shields (e.g., Tempe Terra Formation in northeastern Tharsis;Plescia, 1981). Each of these basaltic landscapes includes a diverse range of channels, which may be related to volcanism, bedrock erosion by overland flows of water, or a combination of both processes. However, distinguishing among these hypotheses requires determining if the channels are primarily features associated with lava flow emplacement, or if they are products of subsequent erosional processes. This can be achieved by considering the morphological characteristics and facies relationships associated with lava flows, which generally involve two end-member emplacement mechanisms: channel-fed growth (e.g.,Booth and Self, 1973;Baloga et al., 1995;Kilburn, 1996;Harris and Rowland, 2001;Glaze et al., 2009;Harris et al., 2009) and endogenous growth (e.g.,Walker, 1991;Hon et al., 1994;Keszthelyi and Denlinger, 1996).

6.2. Channel-fed flows

Channel-fed lavas (Figs. 19,20) can form within a range of lava flow types including ‘a’ā, pāhoehoe, and transitional lavas (e.g., blocky, rubbly, slabby, and platy flows). Lava channels are commonly associated with high lava-discharge rates, which favor the transport of lava in open ‘a’ā channels caused by shear-induced disruption of the lava surface (Macdonald, 1953;Pinkerton and Sparks, 1976;Rowland and Walker, 1990). Viscous tearing of the lava surface also enhances the cooling of the underlying molten lava, thereby resulting in a thermally inefficient lava transport system, relative to endogenous flows. With distance from the vent ‘a’ā channel geometries tend to evolve from being narrow and leveed in the proximal regions to wide and nonleveed near the flow front (Lipman and Banks, 1987;Kilburn and Guest, 1993;Cashman et al., 1999). However, as the flow front continues to advance, the stagnated lateral margins of the ‘a’ā can develop into stationary levees that help to focus the continued through flux of lava. As the eruption progresses, the levees may continue to grow as lava episodically overtops the channel banks to form a combination of overbank flows and rubble avalanches. However, as lava-discharge rates wane, the ‘a’ā channel may partially drain to produce a deep topographic depression bounded by levees on either side.

Pāhoehoe flows tend to be associated with lower lava-discharge rates (Rowland and Walker, 1990), which favor the formation of lobes that are bounded by a continuous surface crust (see the discussion of endogenous flows below for more detail). However, in the ventproximal region, or along segments of a lava pathway where local lava-discharge rates are high (e.g., due to topographic constrictions and/or breakouts of stored lava), lava flow velocities may be sufficiently great that a continuous surface is unable to form. At these localities, pāhoehoe and transitional lavas may also form open channels that can subsequently drain to produce topographic depressions.

Channels within ‘a’ā and pāhoehoe lava flows can be expressed as distinctive landforms that are perched above their surroundings by their high-standing margins. However, large eruptions can also produce broad sheet-like flows that inundate the landscape such that they are confined by pre-eruption topography rather than auto-confining by lateral lava levees (Self et al., 1996,1998;Thordarson and Self, 1998). Theseflood lavas may include one or more preferred pathways that drain as lava-discharge rates wane to produce channel-like depressions that appear to incise into an existing plain. In some cases, lava pathways may carve into the pre-eruption landscape through processes of thermal-mechanical erosion (Baloga et al., 1995;Williams et al., 2000,2001a,b,2005), but this process is exceedingly difficult to discern from remote sensing data alone because the apparentexcess depth of the channel may simply reflect the fact that the preferred pathway formed above the lowest point in the initial landscape.

6.3. Endogenous flows

In contrast to channelized flows, endogenous (e.g., pāhoehoe) lava flows tend to be composed of self-similar lobes (Bruno et al., 1994) that transport lava though thermally insulated internal pathways, which may range from to narrow lava tubes to broad sheet-like regions (Self et al., 1998). Endogenous flows grow as new lobes breakout along existing flow margins (Crown and Baloga, 1999). These breakouts expose fluidal lava that quickly cools to develop a rheological gradient consisting of a brittle outer crust, underlying viscoelastic layer, and inner molten core. Once lobes develop sufficient rigidity to retain incoming lava, they can pressurize and inflate as a network of lobes or coalesce to form a lava-rise plateau (Walker, 1991). Lava-rise plateaus generally have a broad sheet-like geometry, but as they thicken by inflation, they can also generate lava-rise pits above topographic highs in the pre-eruption landscape (Walker, 1991). Lava-rise plateaus and lava-rise pits are important diagnostics of inflation because they are large enough to be observable on high-resolution remote sensing data (e.g., MRO HiRISE and Context Camera (CTX) imagery).

6.4. Facies changes

Distinguishing between channelized and endogenous growth mechanisms is vital for inferring lava emplacement dynamics, modeling flow behavior, and understanding the origin of volcanic plain units on Earth, Mars, and other planetary bodies (Self et al., 1998). However, lava transport mechanisms evolve with distance from source and with time such that structures preserved at the surface of a flow may only represent one part of a complex emplacement history. To address this issue, a facies-based approach can prove fruitful for advancing understanding the formation and characteristics of lava channels.

A facies refers to the suite of characteristics (i.e., appearance, composition, etc.) of a rock unit, or stratified, body that reflects its origin and enables the unit to be distinguished from others around it. In the context of lava facies,Kilburn and Guest (1993) described how lava flows can change continuously from the initial emplacement of isothermal sheets to flows with tubes and channels, which commonly evolve from pāhoehoe to ‘a’ā flows.Kilburn and Guest (1993) also explored how combinations of poorly/well-crusted and sheet/channel zones can be used to establish a continuum of facies. These facies reflect the balance between dynamic processes of lava cooling, crustal growth, and surface stability relative to the mode of lava transport in either thermally insulated internal pathways or open channels. Facies change with distance from the source and at any given location with time, which makes facies relationships a much stronger diagnostic of a unit’s volcanic origin than isolated observations.

In the context of Mars (Fig. 21), small sinuous channel systems in volcanic terrains have been particularly contentious, with proposed origins ranging from fluvial processes and mudflows (e.g.,Murray et al., 2010;El Maarry et al., 2012) to lava flow emplacement (e.g.,Hamilton et al., 2010,2011). These channels are hundreds of meters wide, tens of meters deep, and tens to hundreds of kilometers long. Such channels tend to head from fissures, but they are much smaller than outflow channels. These smaller sinuous channel systems are typically located in the Tharsis and Elysium volcanic provinces (localities TV and E,Fig. 9), but they are also distributed in the southern highlands. The channels can have single-stem or multichannel (anabranching) forms with streamlined islands and terraced margins, which are often thought to be diagnostic of bedrock erosion by water. However, examples of similar landforms are also observed within lava flows on Earth (Soule et al., 2004). Unfortunately, in cases where volcanic eruptions have continued for long periods of time, initial lava flow structures, such as anabranching channel morphologies, may be overprinted and erased by subsequent lava flows. For example, as lava-discharge rates wane, endogenous flow units emplaced under lower lava-discharge conditions may overlie channel-fed flows that formed under peak lava-discharge conditions. The time-scales of emplacement for lava flows on Mars are poorly constrained, butJaeger et al. (2007,2008,2010) advocated that some Martian lava flows, such as the Athabasca Valles flood lava (locality I,Fig. 9), may have been emplaced over a relatively short duration, on the order of several weeks. If so, short-lived, but relatively high lava-discharge rate eruptions (e.g., the December 1974 flow on Kīlauea, Hawai’i) may provide valuable insight into the formation of sinuous lava channels on Mars.

7.1. Cataclysmic megaflooding forms and processes on Earth and Mars

7.2. The circum-Chryse outflow channels region

The chaotic source areas and relatedoutflow channels surrounding the Chryse region of Mars (locality C,Figs. 9 and22) have received extensive study for more than 40 years (Sharp, 1973;Baker and Milton, 1974;Carr, 1979;Baker et al., 1991;Clifford and Parker, 2001;Rodriguez et al., 2003,2005a,b,2006b,2007,2011;Bargery and Wilson, 2011;McIntyre et al., 2012). The upper crustal stratigraphy in this region is thought to consist of interbedded volcanic and sedimentary deposits that were mostly emplaced during the Noachian (Rotto and Tanaka, 1995;Scott and Tanaka, 1986;MacKinnon and Tanaka, 1989). These deposits are hypothesized to contain large populations of buried impact craters (Malin and Edgett, 2001;Frey, 2003;Rodriguez et al., 2005b). Regional investigations indicate that aquifers developed in association with buried craters (Malin and Edgett, 2001;Frey, 2003;Rodriguez et al., 2005b) and tectonic fabrics (Rodriguez et al., 2007). The instabilities within these aquifers probably led to the formation of chaotic terrains. Though not well understood, the instabilities have been linked to conditions such as intrusive magmatism into the hydrosphere/cryosphere (e.g.,Rodriguez et al., 2003,2005b;Harrison and Grimm, 2008), explosive dissociation of upper crustal clathrate deposits (e.g.,Komatsu et al., 2000;Rodriguez et al., 2006b;Gainey and Elwood Madden, 2012), and the thermally insolating effect of hydrated salts (Kargel et al., 2007) and porous sediments (Rodriguez et al., 2011).

The regional mapping ofRotto and Tanaka (1995) portrays a generalized history of channel dissection and includes four outflow channel units that distinguish older and younger, higher (shallow) and lower (deep) channel floors. Their mapping suggests that the chaotic terrains and outflow channels mostly developed between the Late Hesperian and the Early Amazonian. However, recent analysis of very high-resolution imagery shows that portions of Simud, Tiu, and Ares Valles experienced major outflow channel flows during the Early and Middle Amazonian, extending to as recent as ~900 Ma (Rodriguez et al., 2015). Other regional hydrogeologic processes, including valley dissection and groundwater upwelling, may have lasted from the Late Noachian until the Amazonian (Andrews-Hanna and Phillips, 2007;Glotch and Rogers, 2007;Fassett and Head, 2008).

Higher levels within the circum-Chryse outflow channels consist of ~20–50-km-wide canyons, the floors of which are marked by prominent ridges and grooves. Their formation has been attributed to catastrophic floods generated by groundwater eruptions along the flanks of chaotic terrains and plateau zones of subsidence (Rodriguez et al., 2006a). Some of these chaotic terrains are enclosed, while others extend over highlands and modify craters and intercrater plains alike. Associated zones of subsidence consist of complex systems of warped and faulted highlands (Rodriguez et al., 2003,2005a,b). In addition, a few of these outflow channels extend from structures produced by dilational (Hanna and Phillips, 2006;Coleman et al., 2007) and contractional (Rodriguez et al., 2007) tectonism.

Lower levels within the circum-Chryse outflow channels consist of much broader troughs, generally a few hundred kilometers in width. Their floors are marked by faint ridges and include widespread knobs, which may consist of large blocks likely transported by debris flows (Tanaka, 1997,1999;Rodriguez et al., 2006b). The debris flow hypothesis is consistent with the finding of rounded, submeter size clasts and imbricated boulders at the Mars Pathfinder landing site (Golombek et al., 1997;Tanaka, 1997,1999). These debris flows might have been triggered during episodes of large-scale collapse within the Ganges Chasma (Rodriguez et al., 2006a) or by discharges from vast paleolakes within Valles Marineris (locality VM,Fig. 22) (Rotto and Tanaka, 1995;Lucchitta et al., 1994;Warner et al., 2013). In addition, glaciers appear to have significantly contributed to the formational history of the circum-Chryse outflow channels (Lucchitta, 1982,2001), and immense amounts of glacial ice may have occupied the Valles Marineris (Mege and Bourgeois, 2011;Gourronc et al., 2014).

7.3. Megaflood generation processes

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