Constraints on the history of open-basin lakes on Marsfrom the composition and timing of volcanic resurfacing
Timothy A. Goudge,1 John F. Mustard,1 James W. Head,1 and Caleb I. Fassett2
Received 30 April 2012; revised 4 September 2012; accepted 16 October 2012; published 11 December 2012.
[1] Abundant evidence exists for valley network-related fluvial activity near theNoachian-Hesperian transition on Mars, and areally significant quantities of volcanicridged plains were emplaced during this period as well. Thus, it is worthwhile to explorethe hypothesis that lava-water interaction occurred on the surface of Mars at this time. Weanalyzed the morphology, physical properties, composition, and surface ages of thirtyopen-basin lakes (topographic lows with both an inlet valley and an outlet valley) thatwere also resurfaced by volcanic flows. Hyperspectral imaging data from the CompactReconnaissance Imaging Spectrometer for Mars (CRISM) and Observatoire pour laMinéralogie, l’Eau, les Glaces et l’Activité (OMEGA) instruments indicate that of the30 basins, 12 exhibit the spectral properties of basaltic lithologies with diagnosticabsorptions of olivine, high-calcium pyroxene and low-calcium pyroxene. Anolivine/high-calcium pyroxene mixture is the most commonly identified mineralassemblage, consistent with other Hesperian-aged volcanic units. Therefore, ourmineralogical results for over a third of the open-basin lake floors analyzed support priorinterpretations that the basins were resurfaced by volcanic flows. Supporting evidence forresurfacing by volcanic flows is also given by the observed morphology and physicalproperties (e.g., surface roughness, thermal inertia) of the resurfaced open-basin lakefloors. In these 30 examples, however, no evidence was found for lava-water interaction.The ages of emplacement, derived through counts of superposed craters, of all 30 of theopen-basin lake volcanic resurfacing units show that resurfacing began in the LateNoachian, near the Noachian-Hesperian boundary, was concentrated in the Hesperian,and continued into the Early Amazonian. The lack of geologic features indicative oflava-water interaction suggests that the basins were likely to have been mostly devoid ofwater at the time of the latest phase of volcanic resurfacing. We conclude that there is nogeologic evidence that suggests the fluvial activity associated with the studied paleolakeswas coeval with the emplacement of the observed volcanic resurfacing units, several ofwhich date to the Late Noachian-Early Hesperian.
Citation: Goudge, T. A., J. F. Mustard, J. W. Head, and C. I. Fassett (2012), Constraints on the history of open-basin lakes onMars from the composition and timing of volcanic resurfacing, J. Geophys. Res., 117, E00J21, doi:10.1029/2012JE004115.
1. Introduction
[2] Extensive volcanic plains units are one of the oldestknown and best-documented features of the Martian surface[e.g., Greeley and Spudis, 1981; Scott and Tanaka, 1986;Greeley and Guest, 1987]. These volcanic units are typically
Late Noachian to Hesperian in age, and are thought to haveresurfaced at least �30% of the Martian surface during thistime period [Scott and Tanaka, 1986; Greeley and Guest,1987; Head et al., 2002]. The emplacement of volcanicplains in the Hesperian directly follows a period earlier inMartian history when fluvial activity was common, recordedby distinct morphologies such as large valley networks andpaleolake basins [e.g., Pieri, 1980; Goldspiel and Squyres,1991; Cabrol and Grin, 1999, 2001; Howard et al., 2005;Irwin et al., 2005; Fassett and Head, 2008a, 2008b]. It isthought that much of the fluvial activity that created suchfeatures ended near the Noachian-Hesperian boundary[Irwin et al., 2005; Fassett and Head, 2008b; Hoke andHynek, 2009; Mangold et al., 2012], and so preserved evi-dence of interaction between fluvial and volcanic activity atthis critical junction in Martian history is an intriguingpossibility.
1Department of Geological Sciences, Brown University, Providence,Rhode Island, USA.
2Department of Astronomy, Mount Holyoke College, South Hadley,Massachusetts, USA.
Corresponding author: T. A. Goudge, Department of GeologicalSciences, Brown University, 324 Brook St., PO Box 1846, Providence,RI 02912, USA. ([email protected])
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[3] Paleolake basins provide a topographic depression thatlava may have ponded in, and many paleolake basins arecontained within ancient impact craters [e.g., Goldspiel andSquyres, 1991; Cabrol and Grin, 1999, 2001; Irwin et al.,2005; Fassett and Head, 2008a; Hauber et al., 2009].Impact craters can enable surface volcanic eruptions throughreductions in the crustal thickness, which may be furtheraided by deeply fractured zones under impact basins [e.g.,Pike, 1971; Schultz, 1976, 1978; Head and Wilson, 1992].Open-basin lakes, defined as hydrological basins with bothan inlet valley and an outlet valley, require that water musthave ponded in the basin to at least the topographic level ofthe outlet valley head before overflowing [e.g., Cabrol andGrin, 1999; Fassett and Head, 2008a].[4] In situ and orbital observations show that volcanic
resurfacing has affected many open-basin lakes [Goldspieland Squyres, 1991; Squyres et al., 2004; Fassett andHead, 2008a; Goudge et al., 2012]. A recent study of themorphology of a catalog of 226 open-basin lakes has shownthat a variety of geologic resurfacing and modifying pro-cesses have played an important role in the post-lacustrineactivity history of open-basin lakes on Mars [Goudge et al.,2012]. From this study, it was shown that the emplacementof volcanic plains units in particular was one of the dominantgeologic processes responsible for post-lacustrine activityresurfacing, with 96 of the 226 (�42%) open-basin lakesexamined classified as volcanically resurfaced based on adistinctive morphology of the basin interiors [Goudge et al.,2012]. In this investigation, we present a detailed study of30 open-basin lakes classified as volcanically resurfaced byGoudge et al. [2012], to further assess their morphology,physical surface properties (e.g., thermal inertia), composi-tion and age of emplacement. The 30 basins chosen for thisanalysis are broadly geographically distributed across theMartian surface (Figure 1 and Table 1), and have relatively
large surface areas, allowing for improved crater countingstatistics compared to open-basin lakes with smaller floorareas.[5] The major goals in this study are: (1) to document
evidence of volcanic resurfacing of these open-basin lakesusing multiple remotely sensed data sets, (2) to help furtherunderstand the timing and emplacement of the resurfacingunits within paleolake basins, and (3) to assess the possibilityof lava-water interaction at these sites. These analyses aredesigned to provide further insight into the history of open-basin lakes on Mars, and help to illuminate aspects of thehydrological cycle during the Noachian and Hesperianperiods.
2. Data Used and Methods
[6] The basin and resurfacing unit morphology, topogra-phy, surface roughness, thermal inertia, bolometric albedo,composition, and age of emplacement were assessed for30 open-basin lakes identified as volcanically resurfaced byGoudge et al. [2012] based on their morphology. The suite ofdata sets were all geographically referenced and analyzed inESRI’s ArcMap Geographic Information System (GIS)software, which allows for on-the-fly co-registration andmanipulation of diverse data sets.
2.1. Assessment of Basin Resurfacing Unit Morphology
[7] The morphologies of the open-basin lakes and theirassociated resurfacing units were assessed using a combina-tion of �6 m/pixel images from the Context Camera (CTX)instrument aboard the Mars Reconnaissance Orbiter (MRO)spacecraft [Malin et al., 2007], <50 m/pixel images from theHigh Resolution Stereo Camera (HRSC) instrument aboardthe Mars Express (MEx) spacecraft [Neukum et al., 2004],and the 100 m/pixel global mosaic of daytime infrared (IR)
Figure 1. Distribution of all mapped open-basin lakes from Fassett and Head [2008a] and Goudge et al.[2012] (yellow symbols), and the subset of 30 resurfaced open-basin lakes analyzed in this work (blacksymbols). Background is MOLA topography overlain on MOLA hillshade [Smith et al., 2001].
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images from the Thermal Emission Imaging System(THEMIS) instrument aboard the Mars Odyssey spacecraft[Christensen et al., 2004]. Detailed morphology was alsoinvestigated using �0.25 m/pixel images from the HighResolution Imaging Science Experiment (HiRISE) instru-ment aboard the MRO spacecraft [McEwen et al., 2007].[8] The assessment of morphology focused on the recog-
nition of distinctive characteristics indicative of volcanicresurfacing (Figure 2), such as lobate margins and embay-ment relationships, as well as wrinkle ridges, which arecommonly found in Martian volcanic units [e.g.,Greeley andSpudis, 1981; Scott and Tanaka, 1986; Greeley and Guest,1987; Watters, 1991; Head et al., 2002], but are not neces-sarily diagnostic of volcanic resurfacing alone.2.1.1. Morphologies Indicative of Lava-WaterInteraction[9] The image data described above were also examined
closely for features that might indicate lava-water interaction[e.g., Head and Wilson, 2002, 2007; Wilson and Head,2007; Wilson et al., 2012], with focus on: (1) lava deltas,which are formed by lava flowing into standing water[Moore et al., 1973; Mattox and Mangan, 1997; Skilling,2002], (2) littoral cones, which form due to explosive erup-tions of steam as lava flows into a standing body of water[Moore and Ault, 1965; Fisher, 1968; Jurado-Chichay et al.,1996], (3) rootless cones, which form as lava flows ontowater-saturated sediment [Thorarinsson, 1953; Greeley andFagents, 2001], (4) tuyas, which form as lava erupts undera glacial ice sheet [Mathews, 1947; Jones, 1968; Smellie,
2007; Jakobsson and Gudmundsson, 2008], and (5) maarcraters, which form due to phreatomagmatic eruptions fromthe interaction of rising magma and groundwater or surfacewater [Lorenz, 1973, 1986; Gutmann, 1976; White, 1989].All of these feature classes should be large enough to beresolved in the images examined for this study.[10] Lava deltas are formed as erupting lava flows into
standing water, typically the ocean, which causes the rapidcooling of the lava and the creation of a depositional formthat is roughly similar in morphology to a lacustrine delta oralluvial fan [Moore et al., 1973;Mattox and Mangan, 1997];however, the morphology of many lava deltas is influencedby the fact that they are commonly fed by subsurface lavatubes as opposed to incised channels on the topset of thedelta, allowing them to be distinguished from deltas depositedthrough fluvial activity [Mattox and Mangan, 1997; Skilling,2002; Leeder, 2011]. Additionally, lava deltas are prone tothe failure or collapse of the delta toe, making them evenmore distinguishable from fluvial deltas [Mattox and Mangan,1997; Skilling, 2002].[11] Another possible geologic consequence of lava flow-
ing into standing water in the open-basin lakes is the forma-tion of littoral cones. In addition to the creation of a lava delta[Moore et al., 1973; Mattox and Mangan, 1997], such ascenario can result in the rapid heating of the surroundingwater. If the water surrounding the lava flow is vaporizedquickly enough, the explosive expansion of the water vaporwill fragment the lava flow, forming a cone-shaped depositcomposed of pyroclastic material [Moore and Ault, 1965;
Table 1. Locations of Analyzed Resurfaced Open-Basin Lakesa
Basin # Longitude (E) Latitude (N) References
1 �174.86 �14.63 Forsythe and Zimbelman [1995] and Goudge et al. [2012]2 152.75 �11.53 Irwin et al. [2007] and Goudge et al. [2012]3 60.94 21.10 Fassett and Head [2008a] and Goudge et al. [2012]4 �12.32 �21.67 Goldspiel and Squyres [1991] and Goudge et al. [2012]5 �8.59 25.57 Fassett and Head [2008a] and Goudge et al. [2012]6 �7.21 �8.84 Fassett and Head [2008a] and Goudge et al. [2012]7 2.77 �10.63 Fassett and Head [2008a] and Goudge et al. [2012]8 84.96 �2.42 Fassett and Head [2008a] and Goudge et al. [2012]9 89.71 �0.09 Cabrol and Grin [1999, 2001] and Goudge et al. [2012]10 110.86 �2.71 Cabrol and Grin [1999, 2001] and Goudge et al. [2012]11 �23.53 �23.12 Fassett and Head [2008a] and Goudge et al. [2012]12 102.25 �3.42 Fassett and Head [2008a] and Goudge et al. [2012]13 3.88 �27.90 Fassett and Head [2008a] and Goudge et al. [2012]14 31.72 24.39 Fassett and Head [2008a] and Goudge et al. [2012]15 33.57 16.72 Fassett and Head [2008a] and Goudge et al. [2012]16 77.70 18.38 Fassett and Head [2005] and Goudge et al. [2012]17 175.39 �14.40 Grin and Cabrol [1997], Cabrol and Grin [1999, 2001],
Squyres et al. [2004], and Goudge et al. [2012]18 135.38 �6.91 Fassett and Head [2008a] and Goudge et al. [2012]19 134.92 �9.45 Cabrol and Grin [1999, 2001] and Goudge et al. [2012]20 �165.61 �10.22 Cabrol and Grin [1999, 2001] and Goudge et al. [2012]21 62.17 �2.56 Fassett and Head [2008a] and Goudge et al. [2012]22 �18.29 �26.87 Fassett and Head [2008a] and Goudge et al. [2012]23 171.33 �17.44 Cabrol and Grin [1999, 2001] and Goudge et al. [2012]24 �20.51 �22.45 Goldspiel and Squyres [1991] and Goudge et al. [2012]25 �100.69 �38.57 Mangold and Ansan [2006] and Goudge et al. [2012]26 28.76 �0.03 Cabrol and Grin [1999, 2001] and Goudge et al. [2012]27 �11.15 �26.81 Fassett and Head [2008a] and Goudge et al. [2012]28 �5.25 �21.26 Irwin et al. [2007] and Goudge et al. [2012]29 158.54 �22.89 Fassett and Head [2008a] and Goudge et al. [2012]30 39.74 5.16 Fassett and Head [2008a] and Goudge et al. [2012]
aTable lists basin number (assigned in this study), basin location (north positive latitude and east positive longitude) and appropriatereferences.
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Figure 2. Open-basin lake at�11.53�N, 152.75�E (Basin 2 in Table 1) [Irwin et al., 2007; Goudge et al.,2012] that displays morphologic characteristics suggesting it has been volcanically resurfaced. North is upin both images. (a) Overview of volcanically resurfaced open-basin lake. Basin outline, as defined by aMOLA topographic contour [Fassett and Head, 2008a], is shown in blue. White arrows indicate promi-nent wrinkle ridges. Note the smooth surface texture and high crater retention. Location of Figure 2b isindicated by white box. Image is a mosaic of HRSC nadir image h8425_0000 and CTX imagesB20_017403_1658_XN_14S206W and P07_003703_1682_XN_11S206W overlain on the THEMIS day-time IR global mosaic [Christensen et al., 2004]. (b) Basin perimeter embayment in the southwestern por-tion of the basin shown in Figure 2a. The smooth plains unit contained in the basin is clearly embaying therougher, older terrain exterior to the basin. Image is HRSC nadir image h8425_0000.
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Fisher, 1968; Jurado-Chichay et al., 1996]. Littoral conescan be formed either by lava directly entering a standingbody of water [e.g., Moore and Ault, 1965; Fisher, 1968],or through the growth of the cone structure fed by lava-tubes[e.g., Jurado-Chichay et al., 1996]. Littoral cones are largeconstructional features that are found on Earth in the near-shore regions of islands with active volcanism, particularlyHawai’i, and can also occur in association with lava deltas[Moore and Ault, 1965; Fisher, 1968; Jurado-Chichay et al.,1996]. Littoral cones on Earth are typically hundreds ofmeters in diameter, and tens to up to �100 m in height[Moore and Ault, 1965; Fisher, 1968; Jurado-Chichay et al.,1996].[12] If the basin instead contained only small amounts of
standing water or water in the pore space of saturated orpartially saturated sediment at the time of volcanic resurfa-cing, rootless cones or pseudocraters might have formed inthe volcanic unit. As lava flows onto water-saturated sedi-ment, the excess heat from the lava will cause the pore waterto vaporize, and upon reaching a pressure equal to the tensilestrength of the cooling lava, the water vapor will explodethrough the cooling lava unit, creating a rootless cone orpseudocrater [Thorarinsson, 1953; Greeley and Fagents,2001]. Rootless cones are well studied on Earth, especiallyin Iceland [e.g., Thorarinsson, 1953; Greeley and Fagents,2001; Hamilton et al., 2010a, 2010b], and features similarto these have been inferred on the surface of Mars, rangingin size from �30–1000 m in diameter and >25–60 m inheight [Frey et al., 1979; Greeley and Fagents, 2001;Lanagan et al., 2001; Fagents et al., 2002; Fagents andThordarson, 2007]. Rootless cones also typically occur invery dense clusters [Thorarinsson, 1953; Frey et al., 1979;Hamilton et al., 2010b] and are constructional in nature[Thorarinsson, 1953; Greeley and Fagents, 2001], makingthem easy to distinguish from impact craters. One potentiallimit on the production of pseudocraters is the emplacementof a very thick lava flow unit, which might be able to resistthe vapor pressure of the superheated water, preventing aphreatic eruption; however, it is reasonable to expect that atthe margins of the lava flow (i.e., near the basin rims), thethickness would not be so great so as to prevent pseudocraterformation [Thorarinsson, 1953; Greeley and Fagents,2001].[13] If instead lava reaches the surface from below the
paleolake, a tuya or a maar crater may result. Tuyas will beformed if the rising magma body is able to erupt on thesurface below a large ice sheet. This will create a subglacialvolcanic construct that has steep sides and a flat to slightlydomed top, which is termed either a tuya or a table moun-tain. Tuyas form as a subglacial eruption melts the overlyingice sheet, creating a pool of standing water, which the lavawill erupt into. As the volcanic edifice is built up, it maybreach the top of the ice sheet, allowing for the emplacementof sub-aerial lava flows, thus resulting in the flat top. Thetuya will also build out laterally by the creation of lava deltason its flank with very steep foresets that are a consequenceof the geometry of the circumferential pool of standingwater, resulting in the steep-sided tuya [Mathews, 1947;Jones, 1968; Wilson and Head, 2002; Smellie, 2007;Jakobsson and Gudmundsson, 2008; Wilson et al., 2012].Tuyas are very large features, with lengths of several to tensof kilometers, widths of several kilometers and heights of
several hundred meters [Smellie, 2007]. It has also beenhypothesized that tuyas exist on Mars, having typicallylarger dimensions than terrestrial tuyas [e.g., Chapman andTanaka, 2001; Ghatan and Head, 2002; Smellie, 2007].[14] Alternatively, if the rising magma encounters ground-
water or surface water via draining by fissures, a phreato-magmatic, or sometimes phreatic, eruption may occur,resulting in the formation of a maar crater. On Earth, theseeruptions form low, broad craters, with diameters between�100 and 200 m, and depths from �10–200 m [Lorenz,1973, 1986]. Maar craters are also associated with specificlithofacies, such as ejecta deposits, which form a raised ‘rim’tens to �100 m in height above the ground surface [Lorenz,1973, 1986; Gutmann, 1976; White, 1989]. While maar cra-ters are not readily observable on Mars [e.g., Wilson andHead, 1994], they should theoretically form if risingmagma interacts with groundwater or surface water. How-ever, one complication in identifying maar craters is theirpotential morphologic similarity to impact craters. Althoughthis is the case, the steep maar crater walls typically containdistinct morphologic features such as pyroclastic depositsand lahar flows due to the nature of their formation [Lorenz,1973, 1986; Gutmann, 1976; White, 1989] that cause themto differ in morphology from impact craters. Additionally,maar craters often form in multiple episodes, which can resultin crater morphologies that are not perfectly circular or haveadditional cuspate features around their rims [Lorenz, 1973,1986; Gutmann, 1976; White, 1989], making them furtherdistinct from impact craters.
2.2. Assessment of Basin Resurfacing UnitSurface Properties
[15] Topography of the basin floors was analyzed usingMars Orbiter Laser Altimeter (MOLA) gridded topographyand MOLA-derived slope maps [Smith et al., 2001], as wellas HRSC stereo-derived topography [Neukum et al., 2004;Gwinner et al., 2010]. Additionally, the surface roughness ofthe basin floors and resurfacing units were evaluated usingthe MOLA-derived roughness map of Kreslavsky and Head[2000], which uses variations in topography at three lengthscales (0.6, 2.4 and 19.2 km) to estimate the roughness of theMartian surface.[16] The main physical surface properties of the basin
resurfacing units investigated were the thermal inertia andthe bolometric albedo. The thermal inertia values of thebasin resurfacing units were investigated using the globalthermal inertia map produced by Putzig and Mellon [2007].This map was derived from data from the Thermal EmissionSpectrometer (TES) instrument aboard the Mars GlobalSurveyor spacecraft [Christensen et al., 2001; Putzig andMellon, 2007], and offers a quantitative measure of thecompetence of the surface unit, with high thermal inertiavalues indicating competent surface units, such as bedrock,and low values indicating units of low competence, such asloose granular material or dust [e.g., Putzig et al., 2005;Putzig and Mellon, 2007]. Bolometric albedo values derivedfrom the TES instrument [Christensen et al., 2001] were alsoanalyzed.
2.3. Assessment of Basin Resurfacing Unit Composition
[17] The composition of the basin resurfacing units werestudied using hyperspectral data from both the Compact
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Reconnaissance Imaging Spectrometer for Mars (CRISM)instrument aboard the MRO spacecraft [Murchie et al., 2007]and the Observatoire pour la Minéralogie, l’Eau, les Glaceset l’Activité (OMEGA) instrument aboard the MEx space-craft [Bibring et al., 2004]. These imaging spectrometersprovide spectral reflectance data in the visible to near-infrared(VNIR) region of the spectrum (�0.36–3.9 mm for CRISM,and �0.38–5.1 mm for OMEGA) at spatial resolutions of�18–36 m/pixel (targeted CRISM observations) [Murchieet al., 2007] and �0.3–5 km/pixel (OMEGA) [Bibringet al., 2004], and so offer two unique spatial perspectiveson the composition of the basin resurfacing units.[18] OMEGA data were corrected for photometry, known
instrument artifacts [Bibring et al., 2005; Bellucci et al.,2006] and for atmospheric contributions. The atmosphericgas correction assumes that the contribution of the atmo-sphere and the surface are multiplicative, and that atmo-spheric gas absorptions vary with atmospheric path lengthfollowing a power law [Bibring et al., 1989]. Using thisassumption, spectra taken from the summit of OlympusMons and from a lower portion of the volcano are ratioed tocompute an estimate of the atmospheric transmission spec-trum, which is then removed from each observation to esti-mate the surface reflectance [Bibring et al., 2005; Mustardet al., 2005].[19] The CRISM data were first corrected to I/F by
dividing the radiance data returned from the spacecraft witha solar spectrum [Murchie et al., 2007, 2009]. The I/F datawere then corrected for photometry, and the effects ofatmospheric absorptions were removed through a methodsimilar to that used for the OMEGA data as described above(i.e., through the use of a volcano scan correction) [Mustardet al., 2008; Ehlmann et al., 2009; McGuire et al., 2009].[20] The compositional analyses of the basin floor resur-
facing units were performed using both spectral parametermaps [Pelkey et al., 2007; Salvatore et al., 2010], anddetailed spectral analysis. During the spectral analysis,spectra from basin floors were ratioed to exterior terrain withspectrally bland surface materials in an attempt to removeinstrument noise and to bring out the spectral diversity of thearea of interest [e.g., Mustard et al., 2008; Ehlmann et al.,2009], which in this case is the resurfacing unit on thebasin floor. During the spectral analysis, an emphasis was puton identifying crystal field absorptions in the 1 to 2 mm range,which are caused by electronic crystal field transitions ofoctahedrally coordinated Fe2+ in the mineral structure ofsilicate minerals [Burns, 1993].
2.4. Assessment of Basin Resurfacing UnitEmplacement Ages
[21] Emplacement ages of the basin resurfacing units wereevaluated using counts of superposed craters and models ofcrater production for Mars that allow estimates of both rela-tive ages and absolute model ages [e.g., Tanaka, 1986;Ivanov, 2001; Hartmann and Neukum, 2001; Hartmann,2005; Werner and Tanaka, 2011]. Crater counts were per-formed on CTX, HRSC and/or THEMIS global IR mosaicimages.[22] Superposed craters were mapped using the CraterTools
extension for ArcMap, where a best fit circle for each crateris mapped and the crater diameter is calculated in a localsinusoidal projection, thus offering an accurate assessment
of the crater’s diameter [Kneissl et al., 2011]. Craters wereonly included for model age determination if they exhibit araised rim, a depressed center, and superpose the basinresurfacing unit. All craters that were readily identifiableas secondaries, such as highly clustered craters or cratersoccurring in chains [Oberbeck, 1971; Oberbeck et al.,1975], were excluded from age determinations. The finalcrater counts were analyzed using the software CraterStats,which determines a best fit model age based on a nonlinearleast squares fit to a cumulative crater size-frequency dis-tribution over a given range of crater diameters [Michaeland Neukum, 2010].[23] In order to determine the most accurate model age of
emplacement, small crater diameters were not used in thefinal analysis due to the commonly observed downturn insmall craters on Mars due to crater obliteration [e.g.,Hartmann, 1971; Carr, 1992; Hartmann and Neukum, 2001;Fassett and Head, 2008b; Smith et al., 2008]. For themajority of the basins analyzed in this work, only craterslarger than 1 km in diameter were used in the model agedeterminations; however, a portion of the basins (8; �27%)have very few craters larger than 1 km, and so in an attemptto improve the crater counting statistics, a smaller diameterrange was used for these basins to ensure that at least5 counted craters were used in the model age determination.Using this methodology, the smallest crater diameterincluded in our model age determinations was 600 m. Modelage determinations were made using the Neukum productionfunction reported by Ivanov [2001]. The calculated modelages for the resurfaced open-basin lake floors were alsocompared to the model ages of the end of valley networkactivity from Fassett and Head [2008b] derived from theNeukum production function [Ivanov, 2001].[24] In addition to analyzing model ages, period determi-
nations were calculated from the stratigraphic age boundariesof both Hartmann and Neukum [2001] and Werner andTanaka [2011]. The period determinations were made usingthe derived model ages from the Neukum production func-tion [Ivanov, 2001]. Additionally, period determinationswere made using the cumulative number of mapped cratersgreater than or equal to 1 km in diameter normalized to anarea of 106 km2 (i.e.,N (1)) and the period boundaries definedby both Tanaka [1986] and Werner and Tanaka [2011].
3. Results
3.1. Resurfacing Unit and Basin Floor Morphology
[25] The results of the morphologic survey of the 30 res-urfaced open-basin lakes (Figure 1 and Table 1) continue tosupport the interpretation that all 30 have a distinct mor-phology that is diagnostic of volcanic resurfacing (Figures 2and 3a), consistent with previous studies [Fassett and Head,2008a; Goudge et al., 2012]. This morphology includes:(1) lobate margins that embay both the basin perimeter, oftendefined by crater walls, and older material, such as sedi-mentary units associated with lacustrine activity or centralpeaks/peak rings in basins defined by ancient impact craters(Figure 2b) [Goudge et al., 2012]; (2) wrinkle ridges(Figure 2a, white arrows), which are commonly observed onMartian volcanic surfaces [e.g., Watters, 1991]; and (3) highcrater retention, especially at small crater diameters,
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Figure 3. Open-basin lake at �22.89�N, 158.54�E (Basin 29 in Table 1) [Fassett and Head, 2008a;Goudge et al., 2012] displaying physical characteristics indicative of volcanic resurfacing. North is upin all images. (a) THEMIS daytime IR global mosaic [Christensen et al., 2004] showing the smooth floorunit that embays the basin perimeter. (b) MOLA gridded topography [Smith et al., 2001] overlain on theTHEMIS IR global mosaic [Christensen et al., 2004] showing the smooth floor of the basin interior withminimal topographic variation. (c) MOLA-derived roughness map of Kreslavsky and Head [2000] show-ing a dark basin interior, indicative of a smooth surface similar to other volcanic plains units across thesurface of Mars. Image is a false color RGB where the red channel is roughness at a baseline of19.2 km, green is roughness at a baseline of 2.4 km and blue is roughness at a baseline of 0.6 km, anddarker colors indicate lower surface roughness. (d) MOLA-derived slope map [Smith et al., 2001] overlainon the THEMIS IR global mosaic [Christensen et al., 2004] showing a relatively flat, horizontal interiorwith slopes typically <0.5� across much of the basin interior. (e) THEMIS nighttime IR global mosaic[Christensen et al., 2004] showing a relatively warm signature in the basin interior, indicative of morecompetent material than the relatively cooler exterior terrain. (f) TES-derived thermal inertia [Putzigand Mellon, 2007] showing higher thermal inertia in the basin interior, indicative of more competent mate-rial than the exterior terrain. (g) OLINDEX2 parameter map [Salvatore et al., 2010] derived fromOMEGA observation ORB1445_4 overlain on the THEMIS IR global mosaic [Christensen et al., 2004]showing enrichment of olivine in the basin interior compared to the exterior terrain.
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suggesting a competent surface unit [e.g., Fassett and Head,2008a].[26] Based on the survey results, we found no evidence for
the presence of lava deltas [Moore et al., 1973; Mattox andMangan, 1997; Skilling, 2002], littoral cones [Moore andAult, 1965; Fisher, 1968; Jurado-Chichay et al., 1996],rootless cones [Thorarinsson, 1953; Greeley and Fagents,2001], tuyas [Mathews, 1947; Jones, 1968; Smellie, 2007;Jakobsson and Gudmundsson, 2008], or maar craters[Lorenz, 1973, 1986; Gutmann, 1976; White, 1989]. All fiveof these features should have been resolvable with the res-olution of images used, suggesting that they never formed inthese locations, or if so, were buried by later lava flooding.Alternative hypotheses for the lack of these features indica-tive of lava-water interaction are discussed in section 4.4.
3.2. Basin Resurfacing Unit Surface Properties
[27] The analysis of several data sets that give estimates ofbasin floor surface properties (e.g., surface roughness, ther-mal inertia) indicates that the open-basin lake resurfacing
units have physical properties that are distinct from theirsurrounding terrain (Figure 3).[28] Based on the analysis of MOLA gridded topography,
MOLA-derived slope maps [Smith et al., 2001], HRSC stereo-derived topography [Neukum et al., 2004; Gwinner et al.,2010] and MOLA-derived roughness [Kreslavsky andHead, 2000], all 30 analyzed basins have smooth, flat-lyinginteriors with minimal topographic variation (Figures 3b–3d). This smooth topographic signature (Figures 3b and 3c) isconsistent with topographic and roughness signatures iden-tified for larger volcanic units on the surface of Mars [e.g.,Kreslavsky and Head, 2000;Head et al., 2002;Hiesinger andHead, 2004], and the relatively low slopes of the basininteriors (Figure 3d), typically <0.5�, suggest that the interiorresurfacing units were deposited by a mechanism that resultsin a gravitational equipotential surface, such as a fluid lavaflow. Additionally, all of the open-basin lakes analyzed rep-resent distinct topographic lows in comparison to the exteriorterrain (Figure 3b). This is in accordance with how the basinswere originally defined by Fassett and Head [2008a], where
Figure 3. (continued)
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
8 of 24
the authors required that “the observed feature must remain abasin on the basis of its present topography.” This distinctionis important when considering alternate hypotheses for theformation of the valley networks associated with the basinsand the smooth, volcanic floor units, as have been discussedby previous workers [e.g., Leverington and Maxwell, 2004;Leverington, 2006].[29] The thermal properties of the basin floor units suggest
that the basin resurfacing materials have high values ofthermal inertia compared to their surrounding terrain(Figures 3e and 3f). This is indicated in both THEMISnighttime IR data (Figure 3e) [Christensen et al., 2004] andTES-derived thermal inertia (Figure 3f) [Putzig and Mellon,2007]. In the THEMIS nighttime IR data, the basin resur-facing materials all have a relatively ‘warm’ signature,indicating that they retain more heat at night than the sur-rounding terrain [Christensen et al., 2004], which is expec-ted of competent material.
[30] This observation can be quantified using thermalinertia, which shows that the basin interiors have highervalues of thermal inertia than the surrounding terrain(Figure 3f), again suggestive of a competent resurfacing unit[Putzig et al., 2005; Putzig and Mellon, 2007]. Of the30 basins, 26 have a nighttime IR and thermal inertia signa-ture that is warmer/higher than the exterior terrain. The fouroutliers are all located in Arabia Terra, a region of Mars thatis known to be very dusty [e.g., Tanaka, 2000; Ruff andChristensen, 2002; Arvidson et al., 2003]. These fourbasins do show many of the same morphologic indications ofvolcanic resurfacing observed in the other studied basins(Figure 2), so it is possible that there is a thin surficial dustlayer that is overlying the volcanic resurfacing unit, resultingin lower thermal inertia values. Looking at the distribution ofthermal inertia values for only the 26 basins with high ther-mal inertia interiors, the frequency distribution is skewedtoward the higher end of the global data set (Figure 4a),with a mean value of 264.77 (standard deviation = 72.58)J m�2 K�1 s�1/2. This shows the relatively high thermalinertia values for these units, consistent with a mixture ofsand, rocks, boulders, and bedrock, and also consonant withthe presence of a competent volcanic unit in the basin inter-iors below some surficial layer of granular material [Putziget al., 2005; Putzig and Mellon, 2007].[31] For the 26 basins with high thermal inertia, the fre-
quency distribution of bolometric albedo values is skewedtoward the lower end of the global data set (Figure 4b), withan average value of 0.17 (standard deviation = 0.04)[Christensen et al., 2001]. This average bolometric albedowas coupled with the average thermal inertia value to obtain aunit competence classification based on the work of Putziget al. [2005]. The basin resurfacing units fall into the cate-gory ofUnit B from Putzig et al. [2005], which they interpretas “sand, rocks and bedrock; some duricrust.” The majorityof the Syrtis Major and Hesperia Planum regions on Mars,both of which are large expanses of Hesperian-aged volcanicplains [Greeley and Spudis, 1981; Scott and Tanaka, 1986;Greeley and Guest, 1987], are also classified as Unit B byPutzig et al. [2005]. The physical properties of the resurfa-cing units within the studied basins match those of othervolcanic plains units, and are consistent with volcanicresurfacing.
3.3. Basin Resurfacing Unit Composition
[32] Based on the analysis of CRISM and OMEGA VNIRspectral reflectance data, 12 of the studied resurfaced open-basin lakes (40%) have clear spectral signatures associatedwith the resurfacing units in their interior, all of which indicatean enhancement of mafic minerals in their interior compared tothe surrounding terrain (Figure 3g and Table 2). These 12basins have spectral signatures that exhibit diagnostic crystalfield absorptions caused by electronic crystal field transitionsof octahedrally coordinated Fe2+ in the structure of silicateminerals [Burns, 1993] (Figure 5 and Table 3).[33] Specifically, we have identified the presence of the
mafic minerals olivine, low-calcium pyroxene (LCP) andhigh-calcium pyroxene (HCP) (Figures 5 and 6). Olivine isidentified by the presence of a broad, complex absorptioncentered near 1 mm, with the exact band center and shapevarying with changes in Fe content in the mineral structure[King and Ridley, 1987]. Pyroxene is identified by two
Figure 4. (a) Frequency distribution of TES-derivedthermal inertia values [Putzig and Mellon, 2007] for the26 volcanically resurfaced open-basin lakes with high ther-mal inertia interiors (blue curve) compared to the globaldata set (gray curve). Note that the volcanically resurfacedopen-basin lake curve is skewed toward the higher end ofthe thermal inertia range. (b) Frequency distribution ofTES-derived bolometric albedo values [Christensen et al.,2001] for the 26 volcanically resurfaced open-basin lakeswith high thermal inertia interiors (blue curve) comparedto the global data set (gray curve). Note that the volcani-cally resurfaced open-basin lake curve is skewed towardthe lower end of the albedo range.
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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broad absorptions centered at �1 and 2 mm respectively[Adams, 1974]. In the case of pyroxene minerals, the bandcenter of the �1 and 2 mm absorptions is most closelyrelated to Ca content in the mineral structure, with HCPminerals, such as diopside (MgCaSi2O6) and hedenbergite(FeCaSi2O6), having band centers at longer wavelengths(near 1.05 and 2.3 mm), and LCP minerals, such as enstatite(Mg2Si2O6) and ferrosilite (Fe2Si2O6), having band centersat shorter wavelengths (near 0.9 and 1.8 mm) [Adams, 1974;Klima et al., 2007, 2011].[34] The spectral signatures identified in the basin resur-
facing units show a range of different band shapes and cen-ters, which we interpret to represent a range in relativeproportions of olivine, LCP and HCP, with the most commonspectral signature observed being indicative of a combinationof olivine and HCP (Figures 5 and 6). This mineral assem-blage is consistent with the composition of Hesperian-agedvolcanic plains identified across the surface of Mars [e.g.,Mustard et al., 2005; Baratoux et al., 2007; Rogers andChristensen, 2007; Poulet et al., 2009b; Salvatore et al.,2010; Skok et al., 2010] (Figure 5, bottom plots). Whileother workers have determined relative proportions of themafic minerals present based on identified spectral signatures[e.g., Sunshine et al., 1990; Sunshine and Pieters, 1993;Mustard et al., 1997; Baratoux et al., 2007; Poulet et al.,2009a; Skok et al., 2010], such a detailed analysis isbeyond the scope of this work. An additional observation isthat the spectral signatures of the basin floor units are rela-tively consistent across CRISM and OMEGA data withineach basin (Figure 5), which has implications for composi-tional homogeneity.
3.4. Basin Resurfacing Unit Emplacement Ages
[35] From the analysis of crater counts and cumulativesize-frequency distributions for the 30 basins, model agesand stratigraphic periods were determined that represent theemplacement age of the basin resurfacing units (Figures 7and 8 and Tables 4–6). The model ages and period deter-minations show that the resurfacing of the open-basin lakesoccurred over a wide range in Martian history, beginning atapproximately the Noachian-Hesperian boundary and con-tinuing into the Early Amazonian (Figure 8 and Tables 5 and6). The majority of basins appear to be resurfaced during theHesperian (Figure 8 and Tables 5 and 6), consistent with
large emplacements of volcanic plains during this time,which resurfaced �30% of the planet [Scott and Tanaka,1986; Greeley and Guest, 1987; Head et al., 2002]. Addi-tionally, it is interesting to note that some of the basin res-urfacing ages are similar to the time period when valleynetwork activity on a regional-to-global scale ceased[Fassett and Head, 2008b].[36] The derived ages are consistent with previous model
ages for the basin floors of several open-basin lakes deter-mined through crater counting [e.g., Cabrol et al., 1998;Cabrol and Grin, 2001; Irwin et al., 2002; Fassett andHead, 2008a], although some previous workers have sug-gested that these ages are representative of the termination oflacustrine activity [e.g., Cabrol et al., 1998], as opposed toages of resurfacing, as we conclude here.
4. Discussion
4.1. Definitive Evidence for Post-lacustrineVolcanic Resurfacing
[37] It is clear from these observations of the morphology(Figures 2 and 3a), physical properties (Figure 3), and com-position (Figures 3g and 5) of the resurfacing units in theinteriors of the studied open-basin lakes that they are indeedvolcanic in origin, consistent with previous morphologicstudies of resurfaced open-basin lakes [e.g., Fassett andHead, 2008a; Goudge et al., 2012]. These findings furtherconfirm the importance that resurfacing had in the history ofopen-basin lakes on Mars [Goldspiel and Squyres, 1991;Fassett and Head, 2008a; Goudge et al., 2012], which havebeen devoid of fluvial activity for�3.6–3.7 Gyr [Fassett andHead, 2008a, 2008b]. These results also confirm thehypothesis that volcanic resurfacing of paleolake basins wasa widespread process on the surface of Mars, proposed basedon the analysis of both orbital data [e.g., Goldspiel andSquyres, 1991; Fassett and Head, 2008a; Goudge et al.,2012] and in situ exploration at the Gusev crater paleolake[e.g., Squyres et al., 2004].[38] While the work we have discussed thus far falls into
the paradigm of sustained fluvial activity on the surface ofMars forming the inlet and outlet valley networks that feedthese open-basin lakes [e.g., Goldspiel and Squyres, 1991;Cabrol and Grin, 1999; Craddock and Howard, 2002; Irwinet al., 2005; Fassett and Head, 2008a, 2008b], it has alsobeen suggested by previous workers that some of theobserved valleys and associated open-basins were insteadformed by the flow of low viscosity, effusive lava [e.g.,Leverington and Maxwell, 2004; Leverington, 2006].[39] The strongest argument against the formation of the
valley networks and open-basins through volcanic processesis the observation that these basins remain as topographiclows on the Martian surface [Fassett and Head, 2008a](Figure 3b), a point discussed by Irwin et al. [2005]. In orderfor the outlet channels of these basins to have formed, theliquid that filled the basin interior must have ponded to atleast the topographic level of the outlet valley head beforebreaching the basin perimeter and flowing out on to theexterior terrain, requiring a sustained input of the liquidresponsible for filling the basin [Fassett and Head, 2008a].If flowing lava was the liquid responsible for forming theseinlet and outlet valley networks, and such large volumes ofponding lava did occur within these basins, it would be
Table 2. CRISM and OMEGA Observations for Basins withIdentified Mafic Mineralsa
Basin CRISM Observation ID OMEGA Observation ID
3 FRT0000B7B7 ORB0488_46 No Coverage ORB1194_47 No Coverage ORB0485_210 HRL00018A62 ORB1310_211 No Coverage ORB1348_316 FRT0001FB74 ORB2272_421 HRL0000491E ORB1435_022 FRT0001EC17 ORB0353_124 FRT0000503E ORB2262_227 No Coverage ORB1238_528 FRT0000A4C4 ORB0551_229 No Coverage ORB1445_4
aTable lists the basin number (from Table 1), and the ID numbers of theOMEGA and CRISM (where such data exist) observations used for thespectral identification of mafic minerals within the basin interiors.
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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expected that at least the upper portions of the lava wouldcool by conduction, creating a solid thermal boundary layer,similar to terrestrial lava lakes [Peck and Minakami, 1968;Wright et al., 1968; Peck et al., 1977]. When this boundarylayer solidified, it would have been at approximately thetopographic level of the outlet valley head, which is
typically on the order of hundreds of meters above the cur-rent basin floor (Figure 3b).[40] In order for the basin to remain a topographic
depression, as is observed [Fassett and Head, 2008a], therewould have to have been substantial subsidence of the solid,thermal boundary layer, which would have resulted in dis-tinctive lava deflation morphologies [Wright et al., 1968;
Figure 5. Representative ratioed CRISM (Figure 5a) and OMEGA (Figure 5b) spectra from the interiorsof five different volcanically resurfaced open-basin lakes (middle plots) compared with the spectra of SyrtisMajor volcanic smooth plains material (bottom plots) [Skok et al., 2010]. The CRISM and OMEGA spectraare from the same five basins, and are taken from areas that are as geographically close to one another aspossible. See Table 3 for full image information, as well as spectra (numerator and denominator) locations.Listed basin numbers are from Table 1. (a) Top plot shows example numerator and denominator spectraused for computing a ratioed CRISM spectrum. Numerator is characteristic of the mafic rich, volcanic res-urfacing unit in Basin 21, while denominator is spectrally bland material. Lower two plots show CRISMratioed spectra from five different volcanically resurfaced open-basin lakes (middle plot) compared witha spectrum of Syrtis Major volcanic smooth plains material (bottom plot) [Skok et al., 2010]. (b) Top plotshows example numerator and denominator spectra used for computing a ratioed OMEGA spectrum.Numerator is characteristic of the mafic rich, volcanic resurfacing unit in Basin 21, while denominator isspectrally bland material. Lower two plots show OMEGA ratioed spectra from five different volcanicallyresurfaced open-basin lakes (middle plot) compared with a spectrum of Syrtis Major volcanic smoothplains material (bottom plot) [Skok et al., 2010].
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
11 of 24
Thordarson and Self, 1998] as well as ‘high-lava marks’,neither of which are observed within these basins, althoughexamples can be found elsewhere on Mars [see, e.g., Jaegeret al., 2010]. Furthermore, it is likely that more than just anupper thermal boundary layer of the flowing lava wouldhave solidified within the basins due to the requirement ofsustained flow to form the outlet valley [Fassett and Head,2008a], thus making it even more difficult to explain thetopographic lows within these basins if the inlet valleys andoutlet valley were formed by flowing lava.[41] Stratigraphic relationships within these basins also
show clear embayment of the basin perimeter (Figure 2), aswell as older lacustrine deposits, where they occur(Figure 9), indicating that the basins must have experiencedsome period of fluvial activity prior to volcanic resurfacing.Therefore, while it is known that low viscosity, effusive lavaflows can cause erosion of the underlying terrain on Mars toform sinuous channels [e.g., Williams et al., 2005; Hurwitzet al., 2010], we do not believe that such lava erosion
Figure 6. Laboratory spectra of primary mafic mineralsthat have signatures similar to those derived from CRISMand OMEGA ratioed reflectance spectra from the open-basinlake volcanic resurfacing units (Figure 5). Spectra arefrom the CRISM spectral library [CRISM Science Team,2006]. Diopside (HCP) is shown in red and is sampleLAPP97. Enstatite (LCP) is shown in orange and is sampleC2PE30. Forsterite and fayalite (olivine) are shown in green,and are samples C3PO61 and C3PO59 respectively. Spectraare offset for clarity.
Tab
le3.
CRISM
andOMEGA
Observatio
nNum
bers,P
ixelLocations,and
Windo
wSizes
fortheNum
erator
andDenom
inator
Spectra
Usedto
CalculateRatio
Spectra
Sho
wnin
Figure5a
Basin
#
CRISM
OMEGA
Num
erator
Denom
inator
Num
erator
Denom
inator
Observatio
nID
Latitu
de(N
)Lon
gitude
(E)
Pixel
Windo
wLatitu
de(N
)Lon
gitude
(E)
Pixel
Windo
wObservatio
nID
Latitu
de(N
)Lon
gitude
(E)
Pixel
Windo
wLatitu
de(N
)Lon
gitude
(E)
Pixel
Windo
w
3FRT00
00B7B
719
.560
61.845
25�
2519
.422
61.852
25�
25ORB04
88_
419
.560
61.858
5�
519
.495
67.971
15�
1516
FRT00
01FB74
18.332
77.472
15�
1518
.032
77.513
15�
15ORB22
72_
418
.332
77.483
7�
711
.986
76.850
15�
1521
HRL00
0049
1E�2
.657
61.803
15�
15�2
.713
61.808
15�
15ORB14
35_
0�2
.652
61.802
11�
111.38
761
.270
35�
3522
FRT00
01EC17
�27.29
4�1
8.07
015
�15
�27.25
5�1
8.06
715
�15
ORB03
53_1
�27.29
3�1
8.06
815
�15
�22.70
6�1
8.12
535
�35
24FRT00
0050
3E�2
2.49
0�2
0.35
215
�15
�22.44
4�2
0.34
815
�15
ORB22
62_
2�2
2.48
8�2
0.34
97�
15�2
5.27
8�2
0.40
915
�15
a Tableliststhebasinnu
mber(from
Table1;
also
show
nin
Figure5),the
observationID
numberfortheanalyzed
observations
overthebasin(CRISM
andOMEGA),andthelatitud
e(north
positiv
e),lon
gitude
(east
positiv
e)andpixelwindo
wsize
ofthenu
merator
anddeno
minator
spectraused
forprod
ucingtheratio
edspectrashow
nin
Figure5.
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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Figure 7
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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could have created the observed morphologic features (i.e.,basins with inlet channels and an outlet channel) studied inthis work. This conclusion is also in agreement with previ-ous arguments against a volcanic origin for such open-basins[Irwin et al., 2005; Fassett and Head, 2008a].
4.2. Implications of Volcanic Resurfacing UnitComposition
[42] The analysis of the composition of the volcanic res-urfacing units from CRISM and OMEGA hyperspectral dataindicate that where clear spectral signatures are present, theresurfacing units always appear enriched in mafic mineralscompared to their surrounding terrain (Figure 3g). The mostcommonly observed spectral signature for these units is amixture of olivine and HCP, consistent with previouslyidentified examples of Hesperian-aged volcanic smoothplains, and distinct from LCP rich Noachian-aged units[Mustard et al., 2005; Baratoux et al., 2007; Rogers andChristensen, 2007; Poulet et al., 2009b; Salvatore et al.,2010; Skok et al., 2010] (Figure 5, bottom plots). Thiscomposition also supports the model ages derived fromcrater counting and cumulative size-frequency distributions,which indicate a primarily Hesperian age of emplacementfor these units (Figure 8 and Tables 5 and 6).[43] The spectral analysis of the open-basin lake resurfa-
cing units shows that the CRISM and OMEGA spectralsignatures are very similar when looking at individual basinswith both CRISM and OMEGA coverage (Figure 5). Weensured that the spectra extracted from both CRISM andOMEGA are from the same latitude and longitude, within�0.01� (Table 3), which equates to �600 m on the Martiansurface at the equator; however, the difference in spatialscale between the two instruments, a factor of ≥�15[Bibring et al., 2004; Murchie et al., 2007], precludes theextraction of spectra from precisely the same area. Consid-ering such differences in spatial scales, the most probableexplanation for the correspondence in spectral signaturesfrom the CRISM and OMEGA instruments (Figure 5) is thatthe resurfacing units within each basin are relatively com-positionally homogenous.[44] Previous workers have shown that variations in Fe/
Mg content in olivine [Skok et al., 2012] and HCP to LCP
ratios [e.g., Baratoux et al., 2007; Kanner et al., 2007;Poulet et al., 2008, 2009a, 2009b; Skok et al., 2010] aredetectable through CRISM and OMEGA orbital spectros-copy, and so it would be expected that if such variationswere present in the resurfacing units, they would be resolvedin the analyzed spectra. The fact that such spectral variationsare not observed between CRISM and OMEGA data sug-gests that when the basins were resurfaced, the volcanic unitwithin each basin was emplaced as a compositionallyhomogeneous unit across the entire basin floor.
4.3. Timing of Volcanic Resurfacing
[45] The model ages for the emplacement of the volcanicresurfacing units show that the process of volcanic resurfa-cing began in the Late Noachian, was concentrated in theHesperian, and then continued into the Early Amazonian insome locations (Figure 8 and Tables 5 and 6). These agescorrespond well with the global Martian stratigraphy, withmost of the large volcanic provinces emplaced in theNoachian to Hesperian, followed by localized volcanicactivity in the Amazonian [Greeley and Spudis, 1981; Scottand Tanaka, 1986; Tanaka, 1986; Greeley and Guest,1987; Head et al., 2002]. While the model ages foremplacement of the volcanic resurfacing units appear con-centrated in the Hesperian (Figure 8 and Tables 5 and 6),this does not necessarily imply a peak in volcanic resurfacingat this time compared with earlier in Martian history, as thedated volcanic resurfacing units are the youngest of apotential series of volcanic resurfacing units; however, themodel ages for emplacement do imply that volcanic resur-facing in the Hesperian was more prevalent than in theAmazonian.[46] There does not appear to be a correlation between
geographic location of the basins and the emplacement ageof the volcanic resurfacing units (Figure 10). This suggeststhat the temporal signature of emplacement ages (i.e., withvolcanic resurfacing beginning in the Late Noachian, beingfocused in the Hesperian and continuing into the EarlyAmazonian) is not influenced by particular anomalousregions of volcanically resurfaced open-basin lakes, and israther more reflective of the true temporal signature associ-ated with the volcanic resurfacing of open-basin lakes.
Figure 7. Example count areas and counted craters, and cumulative size-frequency distributions for two different volcani-cally resurfaced open-basin lakes analyzed in this work. North is up in both images. (a) Volcanically resurfaced open-basinlake at �8.84�N and�7.21�E (Basin 6 in Table 1) [Fassett and Head, 2008a; Goudge et al., 2012]. Basin perimeter, definedby a MOLA topographic contour [Fassett and Head, 2008a], is indicated by the green line, count area is indicated by theorange line and counted craters used in the model age determination (D > 1 km) are shown as red circles. Image is THEMISdaytime IR global mosaic [Christensen et al., 2004]. (b) Cumulative size-frequency distribution for craters superposed on thevolcanically resurfaced open-basin lake shown in Figure 7a. Only counted craters used in the model age determination areshown. Model age determination is from the Neukum production function [Ivanov, 2001]. The calculated model age isochronis shown in red, and the Hesperian-Amazonian boundary (3.46 Gyr) and the Noachian-Hesperian boundary (3.74 Gyr) areshown in black [Werner and Tanaka, 2011]. (c) Volcanically resurfaced open-basin lake at �21.26�N and �5.25�E (Basin28 in Table 1) [Fassett and Head, 2008a; Goudge et al., 2012]. Basin perimeter, defined by a MOLA topographic contour[Fassett and Head, 2008a], is indicated by the green line, count area is indicated by the orange line and counted craters usedin model age determination (D > 1 km) are shown as red circles. Image is THEMIS daytime IR global mosaic [Christensenet al., 2004]. (d) Cumulative size-frequency distribution for craters superposed on the volcanically resurfaced open-basin lakeshown in Figure 7c. Only counted craters used in the model age determination are shown. Model age determination is fromthe Neukum production function [Ivanov, 2001]. The calculated model age isochron is shown in red, and the Hesperian-Amazonian boundary (3.46 Gyr) and the Noachian-Hesperian boundary (3.74 Gyr) are shown in black [Werner andTanaka, 2011].
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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[47] The emplacement ages give further evidence that mostof the valley networks that feed these open-basin lakes wereformed by flowing water prior to the volcanic resurfacing ofthe paleolake basin floors, as the valleys themselves wereactive until approximately the Noachian-Hesperian boundary[Irwin et al., 2005; Fassett and Head, 2008b; Hoke andHynek, 2009; Mangold et al., 2012], while the majority ofbasin resurfacing units have Hesperian emplacement ages
(Figure 8 and Tables 5 and 6). This can be readily seen whendirectly plotting the population of valley network cessationmodel ages and volcanic resurfacing unit emplacementmodel ages against one another (Figure 8, bottom plot). Thisplot shows that the population of model ages for emplace-ment of the volcanic resurfacing units is largely younger thanthe population of model ages for valley network activitycessation [Fassett and Head, 2008b]. Although this is the
Figure 8
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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case, there are a small number of basins that have model agesof emplacement that are similar to, or only slightly youngerthan, the ages of valley network cessation [Fassett and Head,2008b] (Figure 8).4.3.1. Relationship Between Ages of Valley Networksand Lacustrine Activity[48] In order to make any inference about the relationship
between the volcanic resurfacing ages for the open-basinlake interiors and the timing of lacustrine activity within
these paleolakes, we have made the assumption that thevalley network ages presented by Fassett and Head [2008b]are representative of this period of lacustrine activity. Wefeel that this is a valid assumption for two primary reasons.The first reason is that 15 of the 30 studied open-basin lakesare in fact fed and drained by valley networks dated byFassett and Head [2008b]. Furthermore, all 30 of the studiedopen-basin lakes are fed and drained by valley networks thatare included in the catalog ofHynek et al. [2010].Hynek et al.
Figure 8. Plot of model ages for emplacement of the volcanic resurfacing units within the studied open-basin lakes. Agesshown here were calculated using the Neukum production function [Ivanov, 2001]. See Table 4 for a complete list of basinages. Period boundaries (green lines) are as defined by Werner and Tanaka [2011], where LN = Late Noachian, EH = EarlyHesperian, LH = Late Hesperian, EA = Early Amazonian and MA = Middle Amazonian. Top plot shows a summary of allcalculated model ages for emplacement of the volcanic resurfacing units for the 30 basins analyzed (blue points). Shadedred area indicates the range of model ages for the cessation of valley network activity [Fassett and Head, 2008b], derivedfrom the Neukum production function [Ivanov, 2001], excluding the four youngest valley networks, which are thought tobe unrelated to open-basin lake activity [Fassett and Head, 2008a, 2008b]. Error bars are indicated in black. Bottom plotshows a direct comparison between the volcanically resurfaced open-basin lakes with the oldest (i.e. >3.4 Ga) model agesof emplacement (blue points) and the model ages for valley network cessation, excluding the four youngest valley net-works, from Fassett and Head [2008b] (red points). Note that while there is some overlap in the two populations of ages,the population of volcanic resurfacing model ages is noticeably younger than the population of valley network cessationages. Error bars are indicated in black.
Table 4. Model Ages of Emplacement for the Analyzed Volcanic Resurfacing Unitsa
aTable lists the basin number (from Table 1), the minimum crater size diameter used in model age calculations (Dmin), the number of cratersused for the model age determinations (i.e., N(D ≥ Dmin)), the model age derived from the Neukum production function [Ivanov, 2001] witherrors, and the N(1) value with error, calculated as √n, both normalized to 106 km2. Model age errors are calculated in CraterStats by convertingthe +/�1-sigma error on the model fit for N(1) (which itself is calculated as 1/√n) into uncertainties in the model age based on the utilizedproduction function [Michael and Neukum, 2010]. Note also that the large errors in many of the N(1) values are due to the small count areas ofthe resurfaced open-basin lake floors.
bN(1) values presented in the table are the measured N(1) values for all basins with Dmin = 1 km. For all basins with Dmin < 1 km, the presented N(1)value is the model N(1) value and calculated 1-sigma error obtained from the best fit model age from CraterStats [Michael and Neukum, 2010].
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[2010] compiled a global map of valley networks across thesurface of Mars, and concluded that the large majority ofthese valley networks formed during the Late Noachian-Early Hesperian, consistent with the work of Fassett andHead [2008b], as well as other works dating the ages ofMartian valley networks [e.g., Irwin et al., 2005; Hoke andHynek, 2009; Mangold et al., 2012]. As the lacustrine
activity within the studied open-basin lakes was the result ofthe fluvial activity within their inlet and outlet valley net-works [Fassett and Head, 2008a], it follows that the period oflacustrine activity within these paleolakes must have beencoeval with the period of fluvial activity within the associatedvalley networks.[49] Second, work by many previous authors directly
dating the timing of fluvial activity associated with Martianvalley networks have all concluded that the primary episodeof fluvial erosion in the cratered highlands of Mars occurredearly in the planet’s history, ceasing by approximately theNoachian-Hesperian boundary [e.g., Irwin et al., 2005;Fassett and Head, 2008b; Hoke and Hynek, 2009; Mangoldet al., 2012]. While these works do not date every individualvalley network across the surface of Mars, they all concludebased on similarities of morphology, degradation state andstratigraphic relationships, that a Noachian-Hesperianboundary cessation age is likely to be representative for themajority of Martian highland valley networks [Irwin et al.,2005; Fassett and Head, 2008b; Hoke and Hynek, 2009;Mangold et al., 2012]. Additionally, the multitude of dif-ferent workers that have reached the same conclusion thatMartian valley network formation largely ceased followingthe Noachian-Hesperian boundary [Irwin et al., 2005;Fassett and Head, 2008b; Hoke and Hynek, 2009; Mangoldet al., 2012] adds confidence to the notion that this age isrepresentative of all Martian highland valley networks. It isimportant to note that the term ‘valley network’ as used inthe context above is distinct from younger and morpholog-ically different fluvial features observed across the surface ofMars, such as midlatitude valleys probably associated withthe melting of ice [e.g., Fassett et al., 2010; Howard andMoore, 2011] or channels on impact ejecta blankets offresh craters [e.g., Morgan and Head, 2009; Mangold,2012].
4.4. Implications of the Lack of Lava-WaterInteraction
[50] While there is some overlap between model ages foremplacement of the volcanic resurfacing units and valleynetwork activity cessation (Figure 8), we have observed nogeologic features that would indicate lava-water interaction,with the primary focus of our attention being on detectingevidence such as remnant lava deltas, littoral cones, pseu-docraters, tuyas, and/or maar craters.[51] Lava deltas form as lava flows enter a body of
standing water, creating a deposit that appears roughlysimilar in morphology to lacustrine deltas from the rapid
Table 5. Period Determinations for Emplacement of the AnalyzedVolcanic Resurfacing Unitsa
Basin #
Neukum Production FunctionModel Age Determination N(1) Determination
Hartmann andNeukum [2001]
Period
Werner andTanaka [2011]
PeriodTanaka [1986]
Period
Werner andTanaka [2011]
Period
1 EA EA EA EA2 LH LH EH EH3 LH LH EA EA4 LH LH LH EA5 EH EH EH EH6 EH LH LH LH7 LH LH LH EA8 EH EH EH EH9 EA EA EA EA10 LH EA LH LH11 EA EA LH EA12 LN LN LN LN13 EH EH EH EH14 LH LH LH LH15 LH LH LH LH16 LH EA LH LH17 LH EA LH LH18 EH EH LN LN19 MA EA EA EA20 LN EH EH EH21 EH LH LH LH22 EH LH LH LH23 LN LN LN LN24 LH LH LH LH25 LH LH LH LH26 LN LN LN LN27 LN EH EH EH28 LH LH EA EA29 EA EA LH EA30 EH EH EH EH
aTable lists the basin number (from Table 1), the period determinationsfrom both the Hartmann and Neukum [2001] and the Werner and Tanaka[2011] absolute age boundaries based on the Neukum production function[Ivanov, 2001] model ages (Table 4), and the period determinations fromthe N(1) values (Table 4) based on the Tanaka [1986] period boundariesand the Werner and Tanaka [2011] period boundaries. For perioddeterminations, LN = Late Noachian, EH = Early Hesperian, LH = LateHesperian, EA = Early Amazonian and MA = Middle Amazonian.
Table 6. Summary of Period Determinations for Emplacement of the Analyzed Volcanic Resurfacing Unitsa
Period
Neukum Production Function Model Age Determination N(1) Determination
aTable lists number of model ages of emplacement for the analyzed volcanic resurfacing units within each stratigraphic period for thefour different methods presented in Table 5.
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cooling of the lava [Moore et al., 1973;Mattox and Mangan,1997; Skilling, 2002]. Littoral cones are cone-shapeddeposits composed of pyroclastic material that form due tothe explosive expansion of vaporized water in the nearshoreregions of areas where lava flows enter standing water[Moore and Ault, 1965; Fisher, 1968; Jurado-Chichay et al.,1996]. Pseudocraters are constructional features that form in
dense clusters due to phreatic eruptions as lava flows ontowater-saturated sediment [Thorarinsson, 1953; Greeley andFagents, 2001; Hamilton et al., 2010b]. Tuyas are essen-tially subglacial volcanoes or volcanic constructs that resultin large mesas with steep sides and a flat to slightly domedtop [Mathews, 1947; Jones, 1968; Smellie, 2007; Jakobssonand Gudmundsson, 2008]. Maar craters form as rising
Figure 9. Volcanically resurfaced open-basin lake in Jezero Crater, 18.38�N, 77.70�E (Basin 16 inTable 1) [Fassett and Head, 2005], showing embayment of the older deltaic deposit by the younger volca-nic resurfacing unit. North is up in both images. (a) Overview of the delta deposit in Jezero Crater showingthe smooth, volcanic plains unit on the crater floor. Location of Figure 9b is indicated by white box. CTXimage P03_002387_1987_XI_18N_282W. (b) A portion of the Jezero Crater delta showing embayment bythe volcanic resurfacing unit. White arrows indicate locations of lava embayment of the delta, where thelava flow has subsequently deflated. HiRISE image PSP_003798_1985.
Figure 10. Geographic distribution of period determinations for the emplacement ages of the 30 volca-nically resurfaced open-basin lakes analyzed in this work. Note the lack of correlation between geographiclocation and emplacement age. Period determinations are those derived from absolute model ages from theNeukum production function [Ivanov, 2001] and the Werner and Tanaka [2011] absolute age boundaries(Table 5). Background is MOLA topography overlain on MOLA hillshade [Smith et al., 2001].
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magma encounters groundwater or surface water that hasdrained to depths via fissures, causing explosive phreato-magmatic eruptions, resulting in low, broad craters [Lorenz,1973, 1986; Gutmann, 1976; White, 1989]. We examinedeach of the resurfaced open-basin lakes for all five of thesegeologic landforms, and found no evidence for intact orremnant features indicative of lava-water interaction.[52] The lack of morphologic evidence for lava-water
interaction of the types detailed above suggests that thebasins were largely devoid of water, either standing orcontained in the pore space of partially saturated sediment, atthe time of resurfacing. If this is the case, it has potentiallyimportant implications for the timing of fluvial activity onMars at the Noachian-Hesperian boundary, and so we haveinvestigated several alternate hypotheses for this observa-tion. The three alternate hypotheses for explaining the lackof features indicative of lava-water interaction investigatedin this section are:[53] 1. The studied basins are not a representative sample
of the volcanically resurfaced open-basin lakes initiallyidentified by Goudge et al. [2012], and the majority of theunstudied basins do contain features indicative of lava-waterinteraction.[54] 2. Features indicative of lava-water interaction that
were initially formed have since been eroded.[55] 3. Features indicative of lava-water interaction that
were initially formed have since been buried by youngerlava flows.4.4.1. Unrepresentative Sample[56] One possible explanation for the lack of observed
features that would indicate lava-water interaction is that thebasins investigated here do not represent an accurate subsetof the volcanically resurfaced open-basin lakes identifiedby Goudge et al. [2012]. In this scenario, the majority ofvolcanically resurfaced open-basin lakes do contain evidenceof lava-water interaction, and we have simply analyzed30 basins without such evidence. We do not believe thisscenario is probable, as the basins analyzed in this study arespread across a wide range of geographic locations (Figure 1and Table 1), have a wide range of model ages (Figure 8 andTable 4), and were chosen to be morphologically represen-tative of the volcanically resurfaced open-basin lakes iden-tified by Goudge et al. [2012]. Therefore, we consider the30 basins to be a representative sample of the previouslyidentified volcanically resurfaced open-basin lakes [Goudgeet al., 2012].4.4.2. Erosion of Evidence of Lava-Water Interaction[57] Another possible explanation for the lack of features
indicating lava-water interaction is that such interaction didin fact occur at these basins, however the geologic evidencehas simply been eroded and obscured or removed. As wehave examined the basins for five separate features that mightindicate lava-water interaction, we examine the possibility oferosion for each geologic feature.[58] Looking at a specific example of a terrestrial lava
delta, lava flow into the Pacific Ocean from Kilauea Volcanoin Hawai’i resulted in a lava delta that was approximately3 km in length and 500 m in width [Mattox and Mangan,1997]. This lava delta size is only slightly smaller thansome of the observed deltas and alluvial fans on Mars [e.g.,Ori et al., 2000; Fassett and Head, 2005; Irwin et al., 2005;Di Achille and Hynek, 2010], suggesting that such features
would have remained preserved on the Martian surface ifthey had formed at the time of volcanic resurfacing. If theopen-basin lakes studied here contained appreciable amountsof standing water at the time of volcanic resurfacing and wereflooded by exterior lava flows, a lava delta should haveformed [Moore et al., 1973;Mattox and Mangan, 1997], andwould have been likely to remain preserved on the surface ofMars; however, no such features are observed, suggestingthat they never formed at these sites.[59] Similar to lava deltas, terrestrial tuyas are very large
features, with lengths ranging from �2–15 km, widthsranging from �1–10 km, and heights ranging from �200–1000 m [Smellie, 2007], and so we expect it is unlikely thatsuch large features would have been eroded. Additionally,features interpreted as tuyas have been observed on Mars[e.g., Chapman and Tanaka, 2001;Ghatan and Head, 2002],and it has been noted that these features have typically largerdimensions than terrestrial tuyas, with widths ranging from�30–120 km and heights ranging from �700–1800 m[Ghatan and Head, 2002]. Furthermore, it was hypothesizedby Ghatan and Head [2002] that the tuyas they studied wereassociated with the emplacement of Early Hesperian-agevolcanic ridged plains. This gives further evidence that suchlarge features should not have been eroded subsequent to thetime of emplacement of the volcanic resurfacing units studiedin this work, and thus should still be visible if they had ini-tially formed.[60] Finally, we consider littoral cones, pseudocraters, and
maar craters together, due to their approximately similarmorphology. Terrestrial littoral cones are large construc-tional features with diameters of hundreds of meters, andtypical heights of 20–80 m, with initial heights of up to100 m [Moore and Ault, 1965; Fisher, 1968; Jurado-Chichay et al., 1996]. Pseudocraters on the Earth are typi-cally tens to hundreds of meters in diameter, while observedexamples on Mars are �30–1000 m in diameter [Greeleyand Fagents, 2001; Fagents et al., 2002; Fagents andThordarson, 2007]. Fagents et al. [2002] have also shownthat pseudocraters on Mars are >25–60 m in height, based onindividual MOLA tracks. Maar craters observed on Earthhave the morphology of low, broad craters, with diametersranging in size from �100–2000 m, depths of �10–200 m,and raised rims, composed of pyroclastic ejecta, of tens to�100 m in height [Lorenz, 1973, 1986].[61] The comparatively small sizes of these three features
makes it more difficult to state with certainty that they wouldremain on the surface if they had been formed in lava unitsin the Early Hesperian; however, their large heights com-pared to rim heights of impact craters of comparable sizes[Pike, 1974; Craddock et al., 1997] make it easier to pre-serve littoral cones, pseudocraters, and maar craters thanimpact craters. Additionally, pseudocraters occur in denseclusters [Thorarinsson, 1953; Frey et al., 1979; Hamiltonet al., 2010b], which would again make it easier to pre-serve some indication of such morphologies than stochasti-cally emplaced, individual impact craters.[62] While it is widely reported that small crater removal
is a significant process acting on the surface of Mars [e.g.,Hartmann, 1971; Carr, 1992; Hartmann and Neukum, 2001;Fassett and Head, 2008b; Smith et al., 2008], Smith et al.[2008] have shown through modeling of crater obliterationprocesses that the time-averaged crater obliteration rate at
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Gusev crater is 4.72 � 2.58 nm a�1, consistent with mea-sured erosion rates in situ [Golombek et al., 2006; Smithet al., 2008]. The volcanically resurfaced Gusev craterpaleolake [Grin and Cabrol, 1997; Cabrol and Grin, 1999,2001; Squyres et al., 2004] is one of the open-basin lakeswe have examined (our Basin 17; Table 1), and so this craterobliteration rate may be a reasonable estimate for thecrater obliteration at the 30 sites analyzed in this work.This crater obliteration rate can then be used to calculate thetotal crater obliteration since the time of emplacement of thevolcanic resurfacing units at each of the open-basin lakesstudied here (Table 7), that is the maximum depth of animpact crater or maar crater, or the total height of a littoralcone or pseudocrater, that could have been completely erasedfrom the surface since the time of emplacement of the vol-canic resurfacing unit [Smith et al., 2008]. It should be notedthat crater obliteration rates account for both the crater ero-sion rate and the crater infilling rate [Smith et al., 2008].[63] The results of the total crater obliteration calculations
(Table 7) show an average obliteration height of 16.37(+1.65/�1.70) m based on the model ages of emplacement
calculated from the Neukum production function [Ivanov,2001], with an absolute maximum crater obliteration heightof 17.94 � 9.81 m. These heights are clearly less than theobserved minimum of 25–60 m for Martian pseudocraterheights [Fagents et al., 2002], and on the very lowest end forthe range of heights of littoral cones [Moore and Ault, 1965;Fisher, 1968; Jurado-Chichay et al., 1996] and depths ofmaar crater [Lorenz, 1973, 1986]. Therefore, it seems prob-able that if pseudocraters, littoral cones or maar craters hadformed at these sites, at least some morphologic evidencewould remain in the basin. This argument is strengthened forpseudocraters and littoral cones by the hypothesis that craterobliteration is primarily achieved by infilling of impactcraters [Hartmann, 1971; Carr, 1992], which would affectthe largely constructional pseudocraters [Thorarinsson,1953; Greeley and Fagents, 2001] and littoral cones[Moore and Ault, 1965; Fisher, 1968; Jurado-Chichay et al.,1996] much less than it does impact craters. While it ispossible that some small-scale features indicative of lava-water interaction were formed at some of these sites andsubsequently eroded, it seems unlikely that post-depositionalerosion can wholly explain the lack of observed geologicfeatures indicative of lava-water interaction.4.4.3. Burial of Evidence of Lava-Water Interactionby Younger Lava Flows[64] One final possible explanation for the lack of observed
features related to lava-water interaction is that the resurfa-cing lava units observed in the open-basin lakes are simplythe youngest flows, and they bury previously existing flowsthat exhibit geologic features caused by lava-water interac-tion. In the case of lava deltas and tuyas, this seems anunlikely explanation due to the large size of such features[e.g., Mattox and Mangan, 1997; Smellie, 2007]. However,a lava flow that is approximately tens to hundreds of metersthick could potentially bury any maar craters, littoral conesor fields of pseudocraters that had formed [Moore and Ault,1965; Fisher, 1968; Lorenz, 1973, 1986; Jurado-Chichayet al., 1996; Fagents et al., 2002]. Such a scenario wherelittoral cones, maar craters, or pseudocraters formed butwere buried with younger lava flows is possible, but testingsuch a hypothesis is very difficult, due to the fact that theevidence for lava-water interaction would be completelyburied if the hypothesis were correct.[65] Irrespective of this possibility, we interpret the visible
volcanic units studied in this work to have been emplacedwhen the paleolake basins were largely free of both standingand pore water based on the lack of surficial features indic-ative of lava-water interaction. Although it is possible thatthese surficial volcanic resurfacing units bury older volcanicresurfacing units that contain evidence for lava-water inter-action, none of the observations presented here support sucha hypothesis.
4.5. Temporal Overlap of Valley Network Activityand Volcanic Resurfacing
[66] In order to assess the null hypothesis that the popu-lations of basin resurfacing ages and ages of cessation ofvalley network activity [Fassett and Head, 2008b] aresample subsets of the same distribution of ages, we per-formed Mann-Whitney U statistical significance tests — acommonly used nonparametric, two-sided statistical signifi-cance test for independent populations [e.g., Wackerly et al.,
Table 7. Calculated Total Crater Obliteration Heightsa
aTable lists the calculated total crater obliteration height at the 30analyzed basin sites since the time of emplacement of the volcanicresurfacing units. The values are calculated using the crater obliterationrate of 4.72 � 2.58 nm a�1 derived from Smith et al. [2008] at the Gusevcrater paleolake, and the model ages presented for each basin in Table 4.Table lists the basin number (from Table 1) and the calculated totalobliteration height with error, in meters, since the time of emplacement ofthe volcanic resurfacing unit, based on the Neukum model ages ofemplacement [Ivanov, 2001]. Also listed are the average and maximumcrater obliteration values, with error, in meters.
GOUDGE ET AL.: VOLCANICALLY RESURFACED OPEN-BASIN LAKES E00J21E00J21
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2002] — on the two populations of ages. Mann-WhitneyU tests were performed using both the entire population ofbasin resurfacing ages and valley network cessation ages,and on a subset of both populations including only thosebasins with model ages for resurfacing >3.6 Ga and only the26 oldest valley network cessation ages of Fassett and Head[2008b], which are those thought to be related to open-basinlake activity [Fassett and Head, 2008a, 2008b]. These testsboth reject the null hypothesis that the basin resurfacing agesand the valley network cessation ages [Fassett and Head,2008b] are subset populations drawn from the same distri-bution of ages at the 95% confidence level.[67] We therefore conclude that these two populations of
ages represent two distinct processes and that their ages donot form part of the same age frequency distribution,although there is some statistical overlap between theemplacement ages for the studied volcanic resurfacing unitsand the timing of the end of global valley network activity[Fassett and Head, 2008b] (Figure 8). The geologic evidenceoutlined above for the lack of features indicative of lava-water interaction supports the hypothesis that the basins werelargely devoid of water at the time of at least the latestepisode of resurfacing. This conclusion suggests that theemplacement of Late Noachian-Hesperian-aged volcanicsmooth plains on the floors of these open-basin lakes was notcoeval with appreciable amounts of lacustrine or fluvialactivity within the paleolake basins, as such a co-occurrencewould be likely to have been recorded in the geologic recordin one or more of the ways described above. This conclusionsupports a climate scenario whereby there was at least somephase of drying at these open-basin lake sites following theend of the valley network activity associated with the indi-vidual basins, but prior to at least the surficial episode ofvolcanic resurfacing.
5. Conclusions
[68] Open-Basin Lake Resurfacing: Detailed analysis ofthe morphology, physical properties and mineral composi-tion of the resurfacing units on the floors of 30 open-basinlakes indicates that volcanic resurfacing is the processresponsible for the emplacement of the smooth floor units.These results are consistent with previous morphologicstudies [Goldspiel and Squyres, 1991; Fassett and Head,2008a; Goudge et al., 2012], as well as detailed in situanalyses at the Gusev crater paleolake [e.g., Squyres et al.,2004], that indicate volcanic resurfacing of open-basinlakes has been an important process in the history of Martianpaleolakes.[69] Additionally, the fact that the basins remain as topo-
graphic lows on the surface of Mars [Fassett and Head,2008a] indicates that it is unlikely these open-basin lakesand their associated valley networks formed via flowing lava,and not flowing water, which has been suggested by previousworkers [e.g., Leverington and Maxwell, 2004; Leverington,2006]. Such a process would have left ‘high-lava marks’ andevidence for hundreds of meters of subsidence of a solidifiedvolcanic thermal boundary layer, neither of which areobserved.[70] Mineral Signatures of Resurfacing Material: The
VNIR reflectance properties of the basin floors measuredwith CRISM and OMEGA hyperspectral data show
absorptions diagnostic of mafic minerals. Varying combi-nations of olivine, high-calcium pyroxene and low-calciumpyroxene are identified based on crystal field absorptions inthe 1–2 mm region [King and Ridley, 1987; Adams, 1974;Burns, 1993]. A composition dominated by olivine andhigh-calcium pyroxene is the most commonly identifiedamong the basins. This mineral assemblage is consistentwith the composition of previously identified Hesperian-aged volcanic smooth plains, and distinct from Noachian-aged units, which typically display absorptions indicative oflow-calcium pyroxene rich material [Mustard et al., 2005;Baratoux et al., 2007; Rogers and Christensen, 2007; Pouletet al., 2009b; Salvatore et al., 2010; Skok et al., 2010].Additionally, the correspondence of reflectance propertieswithin individual basins between CRISM and OMEGA datasuggests that the resurfacing units are largely composition-ally homogeneous.[71] Resurfacing Ages and the Transition from Valley
Network Activity to Smooth Plains: Model ages ofemplacement for the volcanic resurfacing units indicate thatthe resurfacing of the open-basin lakes began near theNoachian-Hesperian boundary, was concentrated in theHesperian, and continued into the Early Amazonian. Severalmodel ages for basin volcanic resurfacing exhibit statisticaloverlap with ages of cessation for global valley networkactivity [Fassett and Head, 2008b]. Despite this overlap, wefound no geologic evidence for lava-water interaction atthese sites, supporting the interpretation that the basins werelargely devoid of both surface water and water in the porespace of surficial sediments at the time of at least the latestepisode of volcanic resurfacing. This conclusion implies thatthe volcanic resurfacing of these specific paleolake basinswas not contemporaneous with the fluvial activity thatcarved the inlet and outlet valley networks and filled thebasins with water, and that there must have been someperiod of desiccation at these sites subsequent to the end ofvalley network activity but prior to volcanic resurfacing.
[72] Acknowledgments. We are very grateful for the fine work of theNASA MRO project team and the excellent job by the CRISM ScienceOperations Center (SOC). This work was supported by NASA through asubcontract from the Applied Physics Lab at Johns Hopkins University toJFM. We also gratefully acknowledge support from the Mars Data AnalysisProgram (MDAP-NNX09A146G) and the NASA-ESA Mars Express HighResolution Camera (HRSC) Team activity (JPL 1237163), both to JWH.We thank David Baratoux, Jacob Bleacher and an anonymous reviewerfor detailed and thoughtful reviews and comments that substantiallyimproved the quality of the manuscript. Special thanks are also extendedto J. L. Dickson and W. J. Fripp for help processing image and spectral data,and to B. L. Ehlmann and M. R. Smith for helpful discussions.
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