Pristine impact crater morphology on Pluto – Expectations for New Horizons

Accepted Manuscript Pristine Impact Crater Morphology on Pluto – Expectations for New Horizons Veronica J. Bray, Paul M. Schenk PII: DOI: Reference:

S0019-1035(14)00249-8 http://dx.doi.org/10.1016/j.icarus.2014.05.005 YICAR 11087

To appear in:

Icarus

Received Date: Revised Date: Accepted Date:

13 January 2014 29 April 2014 5 May 2014

Please cite this article as: Bray, V.J., Schenk, P.M., Pristine Impact Crater Morphology on Pluto – Expectations for New Horizons, Icarus (2014), doi: http://dx.doi.org/10.1016/j.icarus.2014.05.005

This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

Pristine Impact Crater Morphology on Pluto – Expectations for New Horizons

Veronica J. Bray1* and Paul M. Schenk2

1

Lunar and Planetary Laboratory, University of Arizona, Tucson, Arizona, 85721, USA 2

Lunar and Planetary Institute, 3600 Bay Area Blvd., Houston, Texas, 77058, USA

*Corresponding Author: Dr. Veronica J. Bray Lunar and Planetary Laboratory, 1541 E. University Blvd., Tucson, AZ 85721 520-626-1967 E-mail: [email protected]

ABSTRACT: This paper combines previous cratering studies and numerical modeling of the impact process at different impact velocities to predict crater morphology on Pluto. As an icy body, Pluto’s craters are expected to be similar in morphology to those on the icy satellites: lesser depth-diameter ratios (d/D), shallower wall slopes and the development of central uplifts in craters of smaller rim-to-rim diameter than craters on rocky bodies of similar gravity. The low impact velocity of the Pluto system (~ 2 km s-1) might cause deviation from this generalization as the simulations presented in this work suggest that decreasing impact velocity from 10 km s-1 to 2 km s-1 results in deeper craters (larger d/D) and a simple-to-complex transition diameter at larger crater sizes than predicted based on gravity scaling alone (D > 6 km). Conversely, decreasing impact velocity from 2 km s-1 to 300 m s-1 produced smaller d/D, akin to the lower d/D noted for secondary craters. This complex relationship between impact velocity and d/D suggests that there might be a larger range of ‘pristine’ simple crater depths on Pluto than on bodies with higher mean impact velocity. The low impact velocities and correspondingly low volumes of impact melt generated at Pluto might prevent the occurrence, or limit the size, of floor-pits if their formation involves impact melt water. The presence, or not, of central floor-pit craters on Pluto will thus provide a valuable test of floor-pit formation theories. The presence of summit-pits or concentric craters on Pluto is plausible and would indicate the presence of layering in the near sub-surface. Palimpsests, multi-ring basins and other crater morphologies associated with high heat flow are not expected and would have important implications for Pluto’s thermal history if observed by New Horizons.

1.0

INTRODUCTION

Asteroid and comet impacts provide a means to investigate the sub-surface structure and composition of solar system bodies as the dimensions of the resultant impact craters depend upon the gravity and crustal properties of the target as well as the size, velocity and composition of the impacting body. The New Horizons mission is due to fly-by the Pluto system in Summer 2015 [Stern 2008] and provides the first opportunity to image the Pluto surface in detail, allowing both the appearance and number of its crater population to be studied for the first time. This paper combines previous cratering studies and numerical modeling of the impact process to predict crater morphology on Pluto based on our current understanding of Pluto’s composition, structure and surrounding impactor population. Pluto is the second largest of the dwarf planets, after Eris, with a mass of 1.31 x 1022 kg and a gravity of 0.65 m s-2. Pluto’s surface gravity is intermediate between the large icy Galilean moons and smaller icy satellites of the saturnian system. Pluto’s craters therefore present a sample of cratering on an intermediate icy body. The predicted impact velocity for Pluto ranges from ~1.2 to 2.8 km s-1 [Zahnle et al., 2003; Dobrovolskis et al., 1997; Dell’Oro et al., 2013], the lower bound being set by the body’s escape velocity. The impactor population at Pluto is predicted to be predominantly cometary, comprised of Kuiper Belt Objects (KBOs) and their fragments [e.g., Weissman and Stern 1994; Durda and Stern 2000; Zahnle et al., 2003; Bierhaus and Dones, this issue]. The impactor population thus has a density close to that of water ice, complicated by porosity and non-water ‘impurities’ [e.g., Brown 2014].

Current bulk density estimates for Pluto itself range from 1830 to 2050 kg m3, consistent with a composition of roughly 50 - 70 % rock and 30 - 50 % water ice by mass [McKinnon et al., 1997; Olkin et al., 2003; Hussmann et al., 2006; Schubert et al., 2010; Robuchon and Nimmo 2011]. This mass is thought to be differentiated into a predominantly water-ice crust above a silicate mantle [e.g., Owen et al., 1993; Robuchon and Nimmo 2011]. Consideration of the impact crater morphology on Pluto thus heavily relies on comparison to craters on other icy bodies of the solar system, especially those with Pluto-like gravities. In addition to water ice, spectroscopic analyses also indicate the presence of volatile ices at the surface including N2, CO, CH4 and CO2 [e.g., Grundy et al., 2002; Grundy et al., 2013]. As with all bodies with icy crusts, the question of whether Pluto’s ice crust melts at depth to form a sub-surface ocean is one of astrobiological importance. The radiogenic heat flow estimated at Pluto based on its predicted volume of radiogenic elements is ~2.7 to 5.4 mW m-2 [Robuchon and Nimmo 2011] - enough to melt its lower water ice crust and maintain a convective ice shell and sub-surface ocean [McKinnon 2006; Hussmann H. et al., 2006]. The possible impact-generation of Charon and its orbital evolution into a mutually synchronous system with Pluto will have provided both impact heat and dissipation of tidal energy. The early heat flow of Pluto is thus expected to be above that derived from radiogenic heating [e.g., Dobrovolski et al., 1997; Collins and Barr, 2008]. The magnitude of this crustal heat flow over recorded geologic time can be investigated through analysis of the surface features. For example, heat flow affects impact crater morphology both through influencing the initial crater shape [e.g., Bray et al., 2014] and

through later alteration of the crater topography [e.g., Dombard and McKinnon 2006; Bland et al., 2012]. Robuchon and Nimmo [2011] modeled a range of crustal compositions and rheologies and find that with certain shell viscosity and ocean potassium contents a present day ocean beneath an ice shell thickness of 165 km is plausible but not certain. Compositional differentiation of the ice shell itself might also be possible due to the different melting depths of N2, CO, CH4 and CO2 ices [Robuchon and Nimmo 2011; Battaglia 2013]. Crustal layering can also form through atmosphere-surface interactions: Young [2013] posits that an upper layer of N2 frost could form, perhaps a meter thick [Bierhaus and Dones, this issue] above the predominantly water-ice bedrock. Any such layering will be evidenced in the morphology of impact craters and other geologic features formed at the surface. Conversely, this ongoing atmosphere-surface interaction might act to erode or infill impact craters, masking fresh crater morphology to some degree and/or infilling the smallest craters.

2.0. APPROACH TO MODELING. Alongside a review of past research into craters on icy bodies, the prediction of crater morphology on Pluto presented in this paper is contributed to by hydrocode modeling results. The hydrocode simulations performed as part of this work utilized the iSALE hydrocode [Collins et al., 2004], a multi-rheology, multi-material expansion of the original SALE code [Amsden et al., 1980] designed to simulate the impact process. Various combinations of impactor size and velocity were performed to produce simple craters on a Pluto-mass body.

The Pluto surface was approximated as homogeneous water ice with a surface temperature of 40 K [e.g., Altenhoff et al., 1988]. The structure and composition of the impactor was simplified to spherical and homogeneous water ice, impacting at an angle of 90 degrees. The departure from the statistically likely 45 degree impact angle is due to the 2D axis-symmetric nature of the iSALE hydrocode. This difference might influence the amount of impact-generated melting for the higher impact velocities, but it not likely to affect the morphology of the simple craters simulated as part of this work. The thermodynamic response of both the projectile and target ice in the simulations was approximated using the Tillotson equation of state for ice Ih [Tillotson, 1962; Ivanov et al., 2002]. The static strength model for ice employed in iSALE was derived from low temperature (77 - 110 °K), high-pressure laboratory data [Durham et al., 1983; Beeman et al., 1988; Rist and Murrell 1994; Weiss and Schulson 1995], and takes account of the material strength dependence on pressure (P), damage and thermal softening. The static strength of damaged and undamaged ice were represented using the Lundborg [1968] approximation:

Y(i,d ) = Y(i,d )0 +

μ(i,d ) P 1+ μ(i,d ) P (Ym −Y(i,d )0 )

(1)

Where Y(i,d)0 is the effective cohesion of the intact (Yi) or damaged (Yd) ice, µ (i,d) are the coefficients of friction of intact (µ i) and damaged (µ d) ice at low pressure, and Ym is the limiting strength at high pressure. Table 1 lists the static ice strength model parameters used in this work [see Collins et al. 2004; 2008, for further parameter descriptions]. To simulate the apparent weakening experienced by the target during impact crater formation, iSALE incorporates the Oscillation Block Model [Ivanov and Kostuchenko

1997] of acoustic fluidization, which supposes that a cratered target comprises of a system of large blocks that oscillate at a constant frequency within a matrix of smaller fragments. Both the blocks and their movement are produced as a result of the passage of the impact shock wave, fracturing and producing seismic shaking in the target. A complete description of the Block Model is presented in Melosh and Ivanov [1999]. The Block Model parameters suitable for impacts into ice are detailed in Bray et al. [2014]. Impactor sizes used for this work were 100, 200, 300 and 400 m in radius producing craters between 0.5 and 20 km in diameter. Impact velocities of 300 m s-1 to 10 km s-1 were used for each impactor size, producing 68 simulated craters. We did not investigate lower velocities as collisions below ~ 100 m s-1 in low temperature ice (assuming a yield strength of 10 MPa) are elastic and not reproduced well by Eulerian codes. The produced crater profiles were measured and are plotted in Figures 1 and 3 to assist in the investigation of the role of impactor velocity in determining final crater dimensions. Given that hydrocode results reflect the programmer-defined inputs for material properties, our lack of understanding of the strength properties of various planetary ices relevant to Pluto prevents the results of these pure water ice simulations from being truly realistic. The results of our simulation work were instead used to inform our predictions of how the low impact velocity on Pluto might affect the crater morphology in a general sense. i.e. a change in simulated crater depth at a particular impact velocity demonstrates the effect of velocity, but the specific value of the crater depth should not be viewed as a prediction itself.

3.0 PREDICTED CRATER MORPHOLOGY ON PLUTO

The final morphology of a pristine impact crater is a result of the collapse of a geometrically simple, bowl-shaped ‘transient crater’. This collapse is controlled by a complex interplay of crustal material strength, target body gravity and impactor energy. Although ice in planetary settings (at temperatures of 10s to 100 Kelvin) is significantly stronger than ice near its melting point [e.g., Durham et al., 1983], it remains inherently weaker than rock. On account of this basic strength difference, the amount of collapse of the transient cavity during crater modification is greater for craters forming in ice than on a rocky body of similar gravity. Extensive work has been completed recording the morphologies and dimensions of craters present throughout the imaged sections of the solar system. These works have revealed fundamental differences in the appearance of craters on rocky and icy bodies [e.g., Passey and Shoemaker 1982; Schenk 1989; 2002; Bray et al., 2008; Barlow 2010; White and Schenk 2011]. Given Pluto’s icy surface, the first, and truly obvious suggestion of this work, is that craters on Pluto will be closer in appearance to those on other icy bodies than those on rocky surfaces. The specifics of this are outlined in the following sections.

3.1. Crater depth and depth/diameter ratios Pristine Crater Depths Observational data for lunar craters has revealed a linear relationship between depth (d) and diameter (D) of simple bowl-shaped craters of the type d = αD where α is a constant of proportionality [Pike 1977]. Simple crater depths on icy bodies also display linear relationships to crater diameter [e.g., Schenk 1989]. Simple craters on small icy satellites

are 30-40% shallower than craters on the larger terrestrial planets [e.g., the saturnian satellites, Schenk 1989], whilst simple crater depths are similar on icy and rocky bodies with comparable gravity [e.g., the Galilean satellites and the Moon, Schenk 2002]. ‘Complex craters’ (those with terraced rims and flat floors or central peaks/pits/peak-rings) are shallower relative to the crater diameter than smaller simple craters on both rocky and icy bodies, resulting in a shallower depth-diameter ratio (d/D) for complex craters [e.g., Pike 1977; Schenk 2002]. This is due to the additional collapse that occurs during the modification of large transient cavities [e.g., Melosh and Ivanov 1999]. The relative weakness of ice results in complex craters on icy bodies being notably shallower, up to 60-70% in the case of Jupiter’s moon Ganymede, than lunar craters of the same diameter [e.g., Croft 1981; Schenk 2002]. However, surface gravity is also a property of planetary bodies that has been shown to be highly influential on d/D. Specifically, higher surface gravities will result in lower depth/diameter ratios of complex craters [Pike 1980]. The effect of impact velocity on simple crater depth-diameter ratio (d/D) has been explored as part of this work (Figure 1). With increasing impact velocity, a projectile will produce wider and deeper craters, as expected on the basis of much previous work [e.g., Melosh 1989 and references therein]. The depth-diameter ratio (d/D, Figure 1B) has a complex progression with increasing impact velocity. Impacts faster than 2 km s-1 lead to smaller d/D ratios as impact velocity increases, in agreement with previous studies [e.g., Cintala 1979; Barnouin et al., 2011]. Conversely, decreasing impact velocity from 2 k m s-1 to 300 m s-1 produced smaller d/D.

This complex relationship between impact

velocity and d/D suggests that there might be a larger range of ‘pristine’ crater depths on

Pluto than on bodies with higher mean impact velocity, with the deepest craters produced by impacts of ~ 2 km s-1. If d/D does have a maximum value for a specific impact velocity and decreases as impact velocity increases/decreases away from this optimum value, it would explain the apparent contradiction between the observations that lower velocity impacts can produce deeper primary craters [e.g., Schultz 1988; Barnouin et al., 2011] and that secondary impacts produce relatively shallow craters [e.g., Bierhaus et al., 2005]. The two-part trend in d/D with impact velocity might be the result of variations in projectile-target coupling: for impact velocities above the sound speed of the target (~ 3 km s-1 in this case), the point-source assumption of crater scaling rules can be applied and the coupling of projectile and target becomes poorer as impact velocity increases [c.f. Baldwin 1963]. For impact velocities below the sound speed of the target, the release of energy from the impact is no longer approximated as a point-source, and standard crater scaling rules no longer apply. Relation to sound speed might explain why the change in trend noted in Figure 1B occurs close to impact velocities of 3 k m s-1. Instead, this work notes this ‘inflection point’ at impact velocities of 2 km s-1. Deviation of the inflection point from the sound speed of the material might be the result of the hydrocode model parameters used for these simulations. For example, an increase in simulated target strength will produce generally deeper craters and increase the diameter at which the simple-to-complex transition occurs, leading to inflection in the d/D trend at different crater diameters. The result of maximum d/D occurring for an impact velocity of 2 km s-1 is thus presented here as an example that a velocity-dependent inflection might occur, not as a firm suggestion of the exact velocity at which maximum d/D would be produced.

Using Pristine Crater Morphology to Study Post-Impact Modification Post-impact modification through erosion/infill and viscous relaxation causes the shallowing of crater depths over time. The rate, and therefore the extent, of viscous relaxation of an impact crater are proportional to the crater diameter and the target body heat flow [Thomas and Squyres 1990]. Comparison between pristine and relaxed crater topography has consequently been used to estimate target body heat flow and rheology [Passey and Shoemaker 1982; Shoemaker et al., 1982]. Likewise, Neish et al. [2013] have used comparison of infilled and eroded Titan craters and their more pristine counterparts on Ganymede to estimate sediment and/or fluvial deposition rates on Titan, assuming the Ganymede craters represent a good analogue for the deepest pristine craters on Titan. Analysis of both pristine and degraded crater depths thus offers a means to study the surface processes on planetary bodies. Recent work by Bray et al. [2014] has shown that the depth of a pristine crater can vary depending on crustal heat flow: higher target heat flow leads to shallower fresh craters, decreases the diameter at which inflections in the depth-diameter (d/D) trend occur and produces central uplifts in craters of smaller rim-to-rim diameter. This presents a complication to using “pristine” crater depths as a reliable starting point for erosion/relaxation estimates. Likewise, the variations of d/D with impact velocity presented in this work (Figure 1B) will further complicate the determination of ‘starting conditions’ (pristine crater d/D) for crater modification models.

Predictions for Pluto: Crater Depth and Depth-Diameter Ratio

Complex craters experience the most collapse, and thus have lower d/D, in weak targets and in high gravity. Pluto’s complex craters are thus expected to be shallower than those on rocky bodies, but not as shallow as complex craters on icy bodies of greater gravity such as the Gailean satellites. Based on its gravity and icy crust, simple crater depth on Pluto is expected to be more akin to that of craters on small icy satellites than craters on rocky bodies (this implies craters 30-40% shallower than simple craters on terrestrial planets [Schenk 1989]). However, Figure 1B shows that depth-diameter ratio (d/D) increases as impact velocity decreases from 10 to 2 km s-1. This suggests that the low mean impact velocity of 2 km s-1 in the Pluto system [c.f. Zanhle et al., 2003] might produce craters deeper than those on other small icy bodies. The change in d/D with impact velocity is not straightforward however, and there are separate d/D-velocity relationships for impact velocities less than and greater than 2 km s-1 (Figure 1C). As 2 km s-1 is the predicted mean impact velocity, it is probable that impacts will also occur at velocities below 2 km s-1, perhaps producing lower d/D craters (Figure 1C) alongside relatively deep craters formed by ~ 2 km s-1 impacts. The complex relationship between impact velocity and crater depth for impacts occurring between 300 m s-1 and 10 km s-1 suggests that there might be a larger range of ‘pristine’ crater depths on Pluto than on bodies with higher mean impact velocity. Knowledge of pristine crater depths on Pluto is important for studies of crater modification over time as the amount of crater infill/erosion/relaxation over time can be inferred quantitatively. The variation of pristine d/D as a function of impact velocity (Figure 1B) and target heat flow at the time of crater formation [Bray et al., 2014] will complicate such estimates as the ‘starting’ depth of the pristine crater will be difficult to

determine. This could mean that changes in d/D due to crater degradation from surfaceatmosphere interactions, viscous relaxation and variations in pristine crater depth are unlikely to be disentangled quantitatively. Investigation of past and present heat flow, crater modification and the possible presence of a sub-surface ocean, on Pluto via cratering studies must then rely on qualitative analysis of crater morphology.

3.2. Collapse of Transient Crater Rims - Wall Slope and Rim Heights The average crater wall slope angle (S) is a useful diagnostic tool when considering the large-scale movement of the target material during impact crater formation as slope angle is a proxy for the ‘effective’ dynamic coefficient of friction (µ eff): the weaker the target, the more extensive the collapse and the shallower the wall slope. For simple craters, wall slope angles increase as crater diameter increases until a stability threshold is met (~30 degrees in the case of the Moon [Pike 1977], ~25 degrees in the case of Ganymede [Bray et al., 2008; Figure 3]. After this maximum slope angle is reached, more wall collapse occurs as crater diameter increases, and slopes decrease. This is associated with the formation of complex craters with wall terraces. Wall collapse and terrace formation is more extreme in larger craters and the average wall slopes of craters on both rocky and icy bodies decrease as crater diameter increases [Pike 1976; Wood 1973; Schenk 1991; Bray et al., 2008]. The degree of slope angle decrease as crater size increases is similar for rocky and icy bodies (Figure 3) indicating that the progression in material weakening, and perhaps the mechanism of this weakening, is the same in icy and rocky bodies [e.g., Schenk 1991].

On bodies with comparable gravity (e.g., the Galilean Satellites and the Moon), crater wall slopes on icy bodies are up to 50% shallower than on rocky bodies [e.g., Schenk 1991; Bray et al., 2008], indicating advanced crater collapse in icy craters, probably in response to the relative weakness of icy targets. However, gravity too plays a role in this: the higher the gravity, the more readily the target collapses and the shallower the wall slopes. Crater wall slopes are thus expected to be steeper on low-mass icy bodies than more massive bodies. For example, craters on the small icy saturnian satellites such as Rhea (g = 0.26 m s-2) have crater wall slopes steeper than those on the larger icy moon of Ganymede (g = 1.43 m s-2) [Schenk 1989, Figure 2], perhaps due to less extensive collapse in low-gravity targets allowing crater wall slopes of up to 35 degrees to remain stable. Final crater rim height is a product of the initially produced rim height and any subsequent collapse that will decrease this initial height. Figures 1B and 1D present final crater rim heights (Hr) from the simple crater simulations performed for this work. Relative rim heights (Hr/D) increase as impact velocity increases from 300 m s-1, reaching a maximum value at velocities of 1 km s-1. Hr/D then decreases as impactor velocity increases, trending to a near-constant rim height for impact velocities 8-10 km. The trend noted for velocities above 1 km s-1 supports the observational work of Cintala [1979] who compared observed crater morphometry on Mercury, the Moon and the moons of Mars and suggested that relative rim heights (Hr/D) for simple craters decrease as impact velocity increases.

For complex craters, the height of crater rims offers indirect evidence of the extent of crater wall collapse [e.g., Melosh 1989; Schenk 1991]. Crater rim heights are thus shorter relative to crater diameter for complex craters, which have experienced substantially more rim collapse than simple craters. This is true for both rocky and icy bodies [e.g. Pike 1977; Schenk 1991]. As the degree of collapse, and thus the rim height, is controlled by both the gravity and crustal strength of the target body, the lower the gravity and the stronger the target, the less the collapse experienced in a crater of a given size. Comparison of icy and rocky bodies of similar gravity [c.f. Ganymede and Moon, Schenk 1991; Bray et al., 2012b] demonstrate that complex crater rim heights on icy bodies are up to 50% shorter than their rocky body counterparts, demonstrating more rim collapse in the weaker ice target.

Wall Slope and Rim Height Predictions Complex crater rim heights and wall slopes on Pluto are expected to be shorter/shallower than on rocky bodies of similar gravity due to the relative weakness of ice facilitating rim collapse on icy bodies. The low impact velocity for the plutonian system might complicate rim height predictions for simple craters however as lower velocity impacts are expected to produce taller simple crater rims [e.g., Cintala 1977; this work]. We tentatively suggest that simple crater rims on Pluto will be taller than those produced by higher velocity impacts on other small icy bodies.

3. 3. Uplift of the Crater floor and Formation of Central Peaks

Central peaks are thought to form via uplift of the transient crater floor during the modification phase of impact crater formation [see Melosh 1989]. There is a positive correlation of central peak diameter height and volume with crater diameter [e.g., Pike 1985; Passey and Shoemaker 1982; Schenk 1991, Bray et al., 2012a]. Craters in icy crusts tend to have larger (broader, taller and more voluminous) central peaks than those on dry rocky bodies of comparable gravity [e.g., Moore et al., 1985; Schenk 1989; Bray et al., 2012b]. This difference is thus attributed to the relative weakness of ice, especially at depth in the target where the target material might be warmer and more mobile. Central uplift size can also be influenced by velocity, producing larger uplifts in craters formed by faster impacts [e.g., Bray 2009]. The crater diameter at which central peaks form is inversely proportional to the target body’s gravity and is also influenced by target strength and volatile content [Melosh 1982]. Figure 4 shows the relationship between simple-to-complex (s-c) transition diameter and planetary gravity for rocky and icy bodies. Although the Moon and the Galilean satellites have similar gravity, the difference in surface composition influences the crater diameter at which the s-c transition occurs: both rocky and icy transition diameter scale inversely with g, but the icy trend is approximately an order of magnitude below the terrestrial planet trend [Chapman and McKinnon 1986]. The simple-to-complex transition diameter might also be influenced by impact velocity, leading to the s-c transition occurring at larger crater diameters for bodies with low impact velocity. This possibility is conjecture at this stage as investigations on this subject are ongoing. Historically, image resolution and coverage have also influenced the recorded s-c transition diameter. Both the lower resolution Voyager results of Schenk

[1989] and the more recent Galileo and Cassini-based results of Schenk [2002] and White et al. [2013] are presented in Figure 4 for illustration of this point. On the Moon, the formation of central peaks is preceded at smaller crater diameters by a decrease in crater rim heights [Pike 1977]. The simple-to-complex transition on the Moon thus proceeds through rim collapse forming wall terraces, to peak formation. On Ganymede, an icy body of similar gravity to the Moon (1.43 m s-2), this order is reversed: central peaks are noted in craters as small as 2 km in diameter [Schenk 2002] and a change in rim height trend is noted at larger crater diameters [~ 9 – 12 km, Schenk 1993; Bray et al., 2012b]. The difference in the order of peak development and rim collapse is either a fundamental difference in the crater modification process between rocky and icy bodies (that floor rebound is the more prominent mechanism of modification/collapse in icy targets rather than rim collapse), or might be the result of measurement bias when assessing craters below 2 km with imagery of ~ 0.4 km/pix resolution in the case of Ganymede from Voyager and Galileo.

Central Peaks and the Simple-to-Complex Transition – Predictions for Pluto Based on its icy crust Pluto would generally be expected to display central peaks with diameters approximately one third of the crater diameter, consistent with other icy bodies. This is viewed as an upper bound to peak size as the low impact velocity for the body might decrease the amount of central uplift occurring relative to some other icy bodies. Pluto’s low gravity, icy crust and low impact velocity are comparable to Triton at which the simple-to-complex transition diameter has been identified as approximately 6 km based on d/D transition [Schenk 1992] and 11 km based on the development of central

peak morphology [Croft et al., 1995]. However, the extremely low crater densities on Triton and limited Voyager coverage limits the degree to which Triton can be used as an analog to this end. The prediction of when Pluto’s central peaks will occur in the crater size-morphology progression is thus based on simple gravity scaling from other saturnian and Galilean satellites and is marked on Figure 4 as a black data point. As the relatively low impact velocity might act to delay the formation of central peaks until larger crater diameters, an s-c transition diameter of ~ 5 km is considered a lower bound. The 0.6 – 0.1 km/pix resolution of the Multicolor Visible Imaging Camera (MVIC) and the Long Range Reconnaissance Imager (LORRI) on board New Horizons [Weaver et al., 2008] are capable of providing high enough resolution data to confidently assess the minimum crater diameter in which central peaks are observed. New Horizons data will also produce a new observational data set for an intermediate-sized icy body, providing the opportunity to investigate the difference in morphology progression between rocky and icy bodies: i.e. does central uplift always occur more readily in icy targets?

3.4. Large and Small Impact Related Pits Classic Central Pits Central peaks are not commonly a well-defined single peak of uplifted material at the crater center. Peak morphology is variable and can be linear, off center, comprised of a cluster of peak-segments, etc. [e.g., Xiao et al., 2014]. In cases where a small depression lies at the center of the peak, the morphology is referred to as ‘summit-pit’ (Figure 5B). Summit-pits occur on both rocky and icy bodies [e.g., Barlow 2010; Xiao et al., 2013].

Theory and modeling work [e.g., Senft and Stewart 2012; Bray et al., 2014] suggest that this crater morphology might be the result of the presence of sub-surface layering (nominally a weaker layer at depth). Bray et al. [2014] suggest that a weak layer at a depth approximately one third of the final crater diameter should result in a summit-pit, if the crater is large enough to form a peak. Summit-pit morphology is seen over a wide range of crater sizes [e.g., Barlow 2010]; wherever there are central peaks present. In craters of increasing diameter on rocky bodies the internal morphology generally changes from central peak to peak-ring. On the Galilean satellites, the transition is instead from peaks to central floor-pits – pits that extend below the minimum level of the surrounding crater floor (Figure 5C) [e.g., Passey and Shoemaker 1982]. This crater morphology is seen on Mars, although peak-ring morphology is also present. Central floor-pits are less ubiquitous in the solar system than summit-pits and are observed predominantly on Ganymede, Callisto and Mars but less commonly on smaller icy bodies such as Titania and Tethys. Theories on their formation are usually more tightly tied to the presence of water ice in the target. Although the exact formation mechanism of central floor-pits remains without full consensus, the idea of Croft [1981], that pit formation involves drainage of impact melt and debris into sub-crater fractures, is gaining more support. A suggested extension of Croft’s melt-drainage formation theory, which can also explain the absence of floor-pit craters on small icy bodies, involves the formation of a peak-ring, which acts to confine impact melt to the crater center. The melt then drains, or is lost as vapor, leaving a central pit and over-printing peak-ring morphology [e.g., Bray et al., 2012b]. This process has been demonstrated as physically possible on both rocky and icy bodies, but much more efficient and likely to form pints in

ice-rich crusts [Elder et al., 2012]. The formation of large ‘floor-pits’ might therefore require: 1) enough impact melt to be produced and 2) the transition from peak to peakring craters to occur. Although melt should be present after impact on small icy bodies (smaller than the Galilean satellites and Titan), no craters large enough to host peak-rings are noted which might explain the rarity of floor-pits on these bodies.

Other Impact-Related Pitting In addition to the very notable summit and floor-pits present in large craters, impactrelated pitting also occurs on smaller scales: a well-studied case of layering in the target affecting crater morphology is that of lunar craters formed in unconsolidated regolith that overlies more resistant mare basalt. Quaide and Oberbeck [1968] found that the exact morphology in such cases is dependent on the thickness of the upper weak layer and includes central mounds, fractured flat floors and concentric craters (Figure 5A). Concentric or ‘terraced’ craters are also noted on Mars [Bramson et al., 2014] and have been used to determine the depth to layers of different composition. The final example of impact-related pitting presented here is that of the smallscale pitting noted in and around fresh craters on Mars (Figure 5D). These pits are noted in ponds of impactite and are thought to occur due to explosive gas release from the meltrich deposit (broken rock fragments, ice and melted rock/ice which are then deposited on top of the crater as fall-out from an ejecta plume [Boyce et al., 2012]). Discovery of this small-scale pitting on Vesta [Denevi et al., 2012] has also led to the suggestion that volatile-rich projectiles can contribute to the volatile content of impact deposits, which

can then de-gas, forming pitting. It is unknown whether this style of pitting occurs on icycrusts (rather than ice-rich rocky crusts) as the resolution available for icy bodies is far below that available from the Mars Reconnaissance Orbiter (MRO) and Dawn missions.

Predictions of Pitting on Pluto The thin but resilient nitrogen and methane atmosphere is not considered significant enough to notably influence the impact process on Pluto by decelerating projectiles or screening out certain projectile sizes as do the thick atmospheres of Titan, Venus, Earth, etc. However, the resultant N2 surface frost, [Young 2005] might be a meter or so deep [Bierhaus and Dones, this issue]. A frost layer above water ice ‘bedrock’ could produce terraced craters like those formed in the upper crusts of the Moon and Mars [c.f. Quaide and Oberbeck 1968; Bramson et al., 2014]. Unfortunately, these features would likely be on the order of 10s of meters across and unlikely to be resolved by New Horizons. However, any thicker layering could produce larger scale versions of the terraced crater in Figure 5A. The ubiquity of summit-pit morphology across both icy and rocky bodies of the solar-system suggests that summit-pit craters are highly likely on Pluto. As summit-pits might be due to sub-surface layering within the upper few kilometers of crust [e.g., Greeley et al., 1982; Bray et al., 2014] any summit-pit craters on Pluto would provide evidence of the presence of sub-surface layering.

Simple extrapolation of gravity-

controlled transitions on Ganymede and other icy bodies [See Moore et al., this issue] predicts that Plutonian craters larger than ~70 km should have central floor-pit

morphologies. However, the low impact velocities predicted for the Pluto system [~1.9 km s-1; Zahnle et al., 2003] would result in less efficient production of impact melt than on bodies with similar gravity and higher impact speeds. This factor might inhibit the formation of central floor-pits on Pluto if floor-pit formation requires the drainage or evaporative loss of impact melt water. The presence, or not, of central floor-pit craters on Pluto will thus provide a valuable test of floor-pit formation theories: is melt required for their formation, or not? Small amounts of melt produced by the low velocity impacts in the Pluto system might contribute to a mix of solid and molten ice fragments and could produce smallscale pitting. On Mars, this pitting generally ranges from the limit of resolution (0.25 m for the High Resolution Imaging Science Experiment onboard MRO) to 300 m in diameter for the largest pit in ~ 30 km diameter crater [Tornabene et al., 2012]. New Horizons is unlikely to resolve most individual pits, but may reveal associated mottling of crater floors (Figure 5F)

3.5. Domes and Multi-Ring Structures In craters between ~ 55 and ~ 180 km in diameter on Ganymede, central pits are partially to extensively filled by sub-circular domed deposits of smooth, high albedo material [e.g. Moore and Malin, 1988; Schenk, 1993]. Central dome craters (Figure 6A) are hypothesized to form as a result of deeply penetrating impacts breaching the brittle ice crust and exposing a relatively fluid layer. Warm sub-surface ice then rises to the surface in the crater center and freezes to form a dome of fresh ice [Schenk, 1993]. Evidence

also suggests that heat flow regulates the occurrence of these types of impact morphologies [Schenk et al., 2004]. Palimpsests are nearly featureless circular regions of high albedo with no discernable crater rim or rim-like features [Passey and Shoemaker, 1982]. Palimpsests are found only on Ganymede and Callisto [Figure 6B and C] and are currently believed to form as a result of highly-fluidized ejecta emplacement following an impact into an ice crust with high heat flow which creates little or no remaining crater topography [Passey and Shoemaker, 1982]. On Ganymede and Callisto, multi-ring basins are limited to the oldest and very largest impacts (D > 500 km equivalent crater diameter, Figure 6C). Ring formation is hypothesized to be due to the collapse of the transient crater when the excavation depth is comparable to the thickness of the ice shell [McKinnon and Melosh, 1980].

The

formation of Valhalla-type multi-ring basins thus requires the presence of a sub-surface ocean or low-viscosity layer. On Europa, Multi-ring structures occur at diameters (D ~ 35 - 40 km) [Figure 6D; Moore et al., 1998; Schenk and Turtle, 2009]. This is significantly smaller than expected for the occurrence of multi-ring basins and is thought to be the result of impact into a possibly thin, high heat flow icy crust.

Predictions of High Heat Flow Crater Morphologies on Pluto Extrapolation of gravity-controlled transitions on the Galilean satellites predicts that craters larger than ~100 km on Pluto should display central dome morphology. Likewise, if palimpsest formation scales with surface gravity, we might expect them to occur at crater diameters of greater than 180-200 km on Pluto. Both are within the range of

potentially observable craters. However, these crater morphologies and multi-ring basins are not expected to be observed by New Horizons unless Pluto’s crustal heat flow has been or is significantly higher than current predictions.

4.0 Conclusions This paper combines previous cratering studies and numerical modeling of the impact process to predict crater morphology on Pluto based on current understanding of Pluto’s composition, structure and surrounding impactor population. The Multicolor Visible Imaging Camera (MVIC) and the Long Range Reconnaissance Imager (LORRI) on board New Horizons will produce observational data for a an intermediate-sized icy body, providing the opportunity to investigate the difference in morphology progression between rocky and icy bodies, and the roles of gravity and impact velocity. Crater morphologies observed in, and dimension trends constructed from, New Horizons data will provide a means to assess the presence of layering in Pluto’s sub-surface. As a small icy body, Pluto’s craters are expected to be similar to those on small icy satellites. However, the low impact velocity of the Pluto system (~ 2km/s) might cause some deviation from this generalization: as impact velocity decreases from 10 km s-1 to 2 km s-1 the d/D and Hr/D of simple craters increases, suggesting that Pluto might display simple craters that are deeper and with taller rims than on other icy bodies of similar gravity. Although this work predicts generally deeper simple craters on Pluto, impacts occurring below the average Pluto impact velocity of ~ 2 km s-1 might produce progressively shallower craters, akin to the shallowing noted for secondary craters. These contrasting trends might complicate the identification of old/young craters based on their

depth-diameter ratios as the depth difference might instead be due to impact velocity variations. Complex craters on Pluto are predicted to be deeper than those of icy bodies with higher gravity. Pluto’s low impact velocity might also delay central peak formation to larger crater sizes than predicted based on gravity scaling alone (D > 5 km). Depending on the mechanism that controls the formation of central floor-pits, they might not be present on Pluto. If formation is connected to the presence of impact melt water, then the low impact velocities might prevent floor-pit formation or limit it to smaller pit sizes due to the relatively low volumes of impact melt produced. The presence, or not, of central floor-pit craters on Pluto will thus provide a valuable test of floor-pit formation theories: is melt required for their formation, or not? Summit-pits and small-scale pitting is however expected. Any summit-pit or terraced/concentric craters on Pluto would provide evidence of the presence of sub-surface layering and provide a means to estimate the thickness of those layers. Central domes, palimpsests and multi-ring basins are not predicted to be observed by New Horizons on the basis of current Pluto heat flow estimates. The occurrence of these crater types would have important implications for Pluto’s thermal evolution.

REFERENCES: Altenhoff, W. J., R. Chini, H. Hein, E. Kreysa, P. G. Mezger, C. Salter, and J. B.Schraml (1988). First radio astronomical estimate of the temperature of Pluto. Astron. Astrophys. 190:L15–L17

Amsden, A.A., Ruppel, H.M., and Hirt, C.W. (1980). SALE: A simplified ALE computer program for fluid flow at all speeds. Los Alamos National Laboratory LA-8095, Los Alamos, NM, 101. Baldwin, R. B. (1963). The measure of the Moon. University of Chicago Press. Barlow, N. G. (2010). Central Pit craters; Observations from Mars and Ganymede and implications for formation models. GSA Special Papers 465:15-27. Barnouin, O.S., C.M. Ernst, J.T. Heinick, S. Sugita, M.J. Cintala, D.A. Crawford and T. Matsui (2011). Experimental results investigating the impact velocity effects on crater growth and transient crater diameter-to-depth ratio. Lunar Planet. Sci.Conf. 42, Abstract # 2258. Battaglia, S. M. (2013). Cycling of Volatiles in Triton’s Icy Lithosphere with Implications for Cantaloupe Terrain on Pluto’s Surface. The Pluto System on the Eve of Exploration by New Horizons: Perspectives and Predictions, Laurel, Maryland. Beeman, M., Durham, W. B., and Kirby, S. H. (1988). Friction of ice, J. Geophys. Res. 93:7625-7633. Bierhaus, E.B., Dones, L. (2014). Craters and Ejecta on Pluto and its Entourage. Icarus, this issue. Bierhaus, E.B., Chapman, C.R. Merline, W.J. (2005). Secondary craters on Europa and implications for cratered surfaces. Nature 437:1125-1127. Bland, M. T., Singer, K. N., McKinnon, W. B., Schenk, P. M. (2012) Enceladus' extreme heat flux as revealed by its relaxed craters. Geophysical Research Letters, 39:17204. Boyce, J. M., L. Wilson, P. J. Mouginis-Mark, C. W. Hamilton, L. L. Tornabene (2012). Origin of small pits in martian impact craters. Icarus 221:262-275.

Bramson A. M., S. Byrne, N. E. Putzig, S. Mattson, J. J. Plaut, J. W. Holt (2014). Thick excess water ice in Arcadia Planitia, Mars. Lunar Planet. Sci.Conf. 45, Abstract #2120. Bray, V. J. (2009). Impact Crater Formation on the Icy Galilean Satellites. PhD Thesis, Imperial College London. Bray, V. J., Collins, G. S., Morgan, J. V., and Schenk, P. M. (2008). The effect of target properties on crater morphology: Comparison of central peak craters on the Moon and Ganymede. Meteoritics and Planet. Sci. 43(12):1979-1992. Bray, V. J., C. Atwood-Stone, and A. S. McEwen (2012a), Investigating the transition from central peak to peak-ring basins using central feature volume measurements from the Global Lunar DTM 100m, Geophys. Res. Lett. 39:L21201. Bray, V. J., Schenk, P. M., H. J. Melosh, Morgan, J. V., and Collins, G. S. (2012b). Ganymede crater dimensions – Implications for central peak and central pit formation and development. Icarus 217:115–129. Bray, V. J., G. S. Collins, J. V. Morgan, H. J. Melosh and P. M. Schenk (2014). Hydrocode simulation of Ganymede and Europa cratering trends – how thick is Europa’s crust? Icarus 231:394-406. Brown, M. E. (2014). The density of mid-sized Kuiper belt object 2002 UX25 and the formation of the dwarf planets. The Astrophysical Journal Letters, In Press. Chapman, C. R., McKinnon, W. B. (1986). Cratering of Planetary Satellites. In: Satellites. J. A. Burns, M. S. Matthews, Eds., University of Arizona, pp. 492 580.

Cintala, M. J. (1979). Mercurian crater rim heights and some interplanetary comparisons. 10th Lunar and Planet. Sci. Conf., Proceedings. Volume 3. (A80-23677 08-91) New York, Pergamon Press, Inc., pp.2635-2650. Collins, G. S., Melosh, H. J. and Ivanov, B. A. (2004). Damage and deformation in numerical impact simulations. Meteoritics and Planet. Sci. 39:217-231. Collins, G. S., Kenkmann, T., Osinski, G. R. and Wunnemann, K. (2008). Mid-sized complex crater formation in mixed crystalline sedimentary targets: Insight from modelling and observation. Meteoritics and Planetary Science. 43(12):195. Croft, S. K. (1981). Cratering on Ganymede and Callisto: comparisons with the terrestrial planets. Lunar Planet. Sci. Conf. 12, pp 187-189. Croft, S.K., Kargel, J.S., Kirk, R.L., Moore, J.M., Schenk, P.M., & Strom, R.G., 1995, in Neptune and Triton, ed. D.P. Cruikshank (Tucson: University of Arizona Press), 879. Dell’Oro, A., A. Campo Bagatin, P. G. Benavidez, R. A. Alemañ (2013). Statistics of encounters in the trans-Neptunian region. Astronomy and Astrophysics 558 A95 DOI: 10.1051/0004-6361/201321461. Denevi, B. W., D. T. Blewett, D. L. Buczkowski, F. Capaccioni, M. T. Capria, M. C. De Sanctis, W. B. Garry, R. W. Gaskell, L. Le Corre, J.-Y. Li, S. Marchi, T. J. McCoy, A. Nathues, D. P. O’Brien, N. E. Petro, C. M. Pieters, F. Preusker, C. A. Raymond, V. Reddy, C. T. Russell, P. Schenk, J. E. C. Scully, J. M. Sunshine, F. Tosi, D. A. Williams and D. Wyrick (2012). Pitted Terrain on Vesta and Implications for the Presence of Volatiles. Science 12:338 (6104),246-249

Dobrovolskis, A. R., Peale, S. J., and Harris, A. W. (1997), Dynamics of the PlutoCharon Binary in Pluto and Charon, S. A. Stern and D. J. Tholen Editors, University of Arizona Press, Tucson, Arizona. pp159-190. Dombard, A. J., and W. B. McKinnon (2006), Elastoviscoplastic relaxation of impact crater topography with application to Ganymede and Callisto, J. Geophys. Res., 111, E01001,doi:10.1029/2005JE002445. Durda, D.D., Stern, S.A. (2000). Collision rates in the present-day Kuiper Belt and Centaur regions: Applications to surface activation and modification on comets, Kuiper Belt objects, Centaurs, and Pluto–Charon. Icarus 145: 220-229. Durham, W. B., Heard, H. C., and Kirby, S. H. (1983). Experimental deformation of polycrystalline H2O ice at high pressure and low temperature: Preliminary results. Proceedings. Lunar planet. Sci. 14th Conf. J. Geophys. Res. 88:B377-B392. Elder, C. M., V. J. Bray, H. J. Melosh (2012). The theoretical plausibility of central pit formation via melt drainage. Icarus 221:831–843. Greeley, R., Fink, J. H., Gault, D. E. and Guest, J. E. (1982). Experimental simulation of impact cratering on icy satellites. In: Satellites of Jupiter. Morrison, D. editor, University of Arizona Press, Tucson, Arizona. pp. 340-378. Grundy, W.M., B. Schmitt, and E. Quirico (2002). The temperature-dependent spectrum of methane ice I between 0.7 and 5 μm and opportunities for near-infrared remote thermometry. Icarus 155,:486-496. Grundy, W. B., C.B. Olkin, L.A. Young, M.W. Buie, E.F. Young (2013). Near-Infrared Spectral Monitoring of Pluto's Ices: Spatial Distribution and Secular Evolution. 223(2):710–721.

Hussmann, H., Sohl, F., Spohn, T. (2006). Subsurface oceans and deep interiors of medium-sized outer planet satellites and large trans-neptunian objects. Icarus 185:258-273. Ivanov B.A., and Kostuchenko V. N. (1997). Block oscillation model for impact crater collapse. Proceedings, Lunar Planet. Sci. 28th Conf. pp 1655. Ivanov, B.A., Langenhorst, F., Deutsch, A., and Hornemann, U. (2002). How strong was impact-induced CO2 degassing in the K/T event? Numerical modeling of shock recovery experiments. In Catastrophic Events and Mass Extinctions: Impact and Beyond, Special paper 356. Geol. Soc. of Am. pp. Boulder, CO. pp. 587-594. Lundborg, N. (1968). Strength of rock-like materials. International Journal of Rock Mechanics and Mining Sciences. 5:427-454. McKinnon, W.B. (2006). On convective instability in the ice I shells of outer solar system bodies, with detailed application to Callisto. Icarus 183,:435-450.

McKinnon, W.B., Melosh, H.J. ,1980. Evolution of planetary lithospheres: Evidence from Multiringed structures on Ganymede and Callisto. Icarus 44, 454-471.

McKinnon, W.B., Simonelli, D.P., Schubert, G. (1997). Composition, internal structure, and thermal evolution of Pluto and Charon. In: Pluto and Charon. Stern, S.A., Tholen, D.J. Editors, Univ. of Arizona Press, Tucson, pp. 295-343. Melosh H. J. (1982). A schematic model of crater modification by gravity. Journal of Geophysical Research 87:371–380.

Melosh, H. J. (1989). Impact cratering: A geologic process. Research supported by NASA. New York, Oxford University Press (Oxford Monographs on Geology and Geophysics, No. 11). Melosh, H. J. and Ivanov, B. A. (1999). Impact crater collapse. Annu. Rev. Earth Planet. Sci. 27:385-415. Moore, J.M., Malin, M.C., 1988. Dome craters on Ganymede. Geophys. Res. Lett. 15, 225-228. Moore J. M., Horner V. M., and Greeley R. (1985). The geomorphology of Rhea. Proceedings, 15th Lunar Planetary Science Conference. Journal of Geophysical Research 90:C785–C795.

Moore, J.M., and 17 colleagues, 1998. Large impact features on Europa: Results from the Galileo nominal mission. Icarus 135, 127-145.

Moore, J. M., A. D. Howard, P. M. Schenk, W. B. McKinnon, R. T. Pappalardo, R. C. Ewing, E. B. Bierhaus, V. J. Bray, J. R. Spencer (2014) Geology before Pluto: preencounter considerations. Icarus (this issue), In Review. Neish. C. D., R. Kirk, R. Lorenz, Bray, V. J., P. Schenk, B. Stiles, E. Turtle, K. Mitchell, A. Hayes (2013) Topography of Craters on Titan. Icarus 223(1):82-90. Quaide, W. L. and V. R. Oberbeck (1968) Thickness determinations of the lunar surface layer from lunar impact craters. Journal of Geophysical Research 73(16): 2156-2202. Ohnaka M. (1995) A shear failure strength law of rock in the brittle-plastic transition regime. Geophysical Research Letters 22:25–28.

Olkin, C. B., O. G. Franz, and L. H. Wasserman 2003. The mass ratio of Charon to Pluto from Hubble Space Telescope Fine Guidance Sensors Observations. Icarus 164, 254259. Owen, T.C., T.L. Roush, D.P. Cruikshank, J.L. Elliot, L.A. Young, C. de Bergh, B. Schmitt, T.R. Geballe, R.H. Brown, and M.J. Bartholomew (1993). Surface ices and atmospheric composition of Pluto. Science 261:745-748. Passey, Q. R. and Shoemaker, E. M. (1982). Craters and basins on Ganymede and Callisto: Morphological indicators of crustal evolution. In Satellites of Jupiter, edited by Morrison, D. Tucson, Univ. of Arizona Press. pp. 340-378. Pike R. J. (1976). Crater dimensions from Apollo data and supplemental sources. The Moon 15:463–477. Pike R. J. (1977). Size-dependence in the shape of fresh impact craters on the moon. In Impact and explosion cratering, edited by Roddy D. J., Pepin R. O., and Merill R. B. New York: Pergamon Press. pp. 489–509. Pike, R.J., (1980). Control of crater morphology by gravity and target type: Mars, Earth, Moon, Proc. Lunar Planet. Sci. Conf. 11, pp. 2159–2189. Pike, R. J. (1985), Some morphologic systematics of complex impact structures, Meteoritics 20:49–68, doi:10.1111/j.1945-5100.1985.tb00846.x. Rist, M. A. and Murrell S. A. F. (1994). Ice triaxial deformation and fracture. J. Glaciology. 40:305-318. Robouchon, G., Nimmo, F. (2011). Thermal evolution of Pluto and implications for surface tectonics and a subsurface ocean. Icarus 216,:426–439.

Schenk, P.M. (1989). Crater formation and modification on the icy satellites of Uranus and Saturn: depth/diameter and central peak occurance, J. Geophys. Res. 94:38133832. Schenk, P.M. (1991). Ganymede and Callisto: Complex crater formation and planetary crusts. J. Geophys. Res. 96:15635-15664. Schenk, P. M. (1992). Volcanism on Triton. Lunar Planet Sci Conf 23. Pp. 1215-1216. Schenk, P. M. (1993). Central pit and done craters: Exposing the interiors of Ganymede and Callisto. J. Geophys. Res. 98:7475-7498. Schenk, P. M. (2002). Thickness constraints on the icy shells of the Galilean satellites from a comparison of crater shapes. Nature 417:419-421.

Schenk, P.M., Turtle, E.P., 2009. Europa's Impact Craters: Probes of the Icy Shell. In Europa, ed. by R.T. Pappalardo, W.B. McKinnon, K. Khurana (Univ. Arizona Press, Tucson,), pp. 181–199.

Schenk, P.M. , Chapman, C.R., Zahnle, K., Moore, J.M., 2004. Ages and Interiors: the Cratering Record of the Galilean Satellites. In: Bagenal, F., et al., (Eds.) Jupiter: The Planet, Satellites & Magnetosphere, Cambridge Univ. Press, New York, pp. 427-456.

Schubert, G., Hussmann, H., Lainey, V., Matson, D.L., McKinnon, W.B., Sohl, F., Sotin, C., Tobie, G., Turrini, D., Van Hoolst, T. (2010). Evolution of icy satellites. Spac. Sci. Rev. 153:447–484. Schultz, P.H. (1988). Cratering on Mercury – A relook. In: Mercury. Vilas, F., Chapman, C.R., Matthews, M.S. Editors, University of Arizona Press, Tuscon, AZ, pp. 274–335.

Senft, L.E., Stewart, S.T. (2011). Modeling the morphological diversity of impact craters on icy satellites. Icarus 214:67–81. Shoemaker, E.M., Lucchitta, B.K., Plescia, J.B., Quyres, S.W., Wilhelms, D.E. (1982). The Geology of Ganymede. In Satelllites of Jupiter, D. Morrison editor, University of Arizona Press, Tucson, pp.435-520 Stern, A. S. (2008). The New Horizons Pluto Kuiper Belt Mission: An overview with historical context. Space Science Reviews 140:3-21. Thomas, P.J., Squyres, S.W. (1990). Formation of crater palimpsests on Ganymede. J. Geophys. Res. 95:19161-19174. Tillotson, J. M. (1962). Metallic equation of state for hypervelocity impact. General Atomic Report GA-3216. Advanced Research Project Agency, San Diego, pp. 141. Tornabene, L. L., G. R. Osinski, A. S. McEwen, J. M. Boyce, V. J. Bray, C. M. Caudill, J. A. Grant, S. Mattson, and P. J. Mouginis-Mark (2012). Widespread crater-related pitted materials on Mars: Further evidence for the role of target volatiles during the impact process. Icarus, 220:348-368. Weaver, H.A. W. C. Gibson, M. B. Tapley, L. A. Young and S. A. Stern (2008). Overview of the New Horizonz science payload. Space Science Reviews 140:75–91. Weiss, J. and Schulson, E. M. (1995). The failure of fresh-water granular ice under multiaxial compressive loading. Acta Metall. Mater. 43:2303-2315. Weissman, P.R. and Stern, S.A. (1994). The impactor flux in the Pluto-Charon system, Icarus 111: 378-386. White, O. L. and P.M. Schenk (2011). Crater shapes on the Saturnian satellites: new measurements using Cassini stereo images. Lunar Planet. Sci. 42, abstract #2283.

White, O.L., P.M. Schenk, A.J. Dombard (2013). Impact basin relaxation on Rhea and Iapetus and relation to past flow. Icarus 223:699-709. Wood C. A. (1973). Central peak heights and crater origins. Icarus 20:503. Xiao, Z. Z. Zeng and G. Komatsu (2014). A global inventory of central pit craters on the Moon: Distribution, morphology and geometry. Icarus 227:195-201. Young, L. A., (2013). Pluto’s Seasons: new predictions for New Horizons. Ap. J. L. 766 L22-L28. Young, L., C. Olkin, T. Ruhland, E. Young, D. French, K. Shoemaker, R. Galvez and B. Gregory

(2005),

Charon

Occultation

Observed

at

SOAR

and

CTIO,

CTIO/CerroTololo Newsletter, 83:28-30. Zahnle, K., Schenk, P.M., Levison, H., Dones, L. (2003). Cratering rates in the outer Solar System. Icarus 163, 263-289.

Table 1: Impactor Properties and Static Strength Parameters Parameter

Value

Cohesion (yield strength at zero pressure), Yi0

10 MPa**

Damaged cohesion, Yd0

0.5 MPa*

Von Mises plastic limit (yield strength at infinite pressure), Ym

0.11 Gpa**

Coefficient of internal friction, µ i

2**

Damaged coefficient of friction, µ d

0.6*

Melt temperature, Tm

273°K

Thermal softening parameter, ξ

1.2***

Density of Impactor and Target Material, ρ

910 kg m-3

Impact Velocity

1-10km s-1

Sources for parameter values: *Fractured ice values from Beeman et al. [1988]. **Intact ice values from Durham et al. [1983]. ***Derived by fitting an Ohnaka [1995] style trend to Durham et al. [1983] data.

Figure 1: Simulation results: comparison of impact velocity and crater depth for the simple craters produced in this work. Rp refers to projectile radius. A) crater depth, B) depth-diameter ratio. These results suggest that lower impact velocities generally produce relatively deeper craters (higher d/D ratio), but that a minimum velocity can be reached below which the depth of craters becomes shallower as velocity decreases still further. Although this inflection point occurs at an impact velocity of 2 km/s in this work, variations in the model parameters could produce an inflection in the d/D trend at a different velocity.

Figure 2: Crater wall slopes on the Moon (rocky, g = 1.63m/s2; white data points from Pike 1977), Ganymede (icy, g = 1.43m/s2; black data points from Bray et al., 2008) and Rhea (icy, g = 0.26 m/s2; grey data cloud for complex craters from Schenk 1989).

Figure 3: A) Simulation results showing the effect of impact velocity on simple crater rim height for 100. 200, 300 and 400m radius projectiles (Rp). Higher velocities produce larger craters and thus larger rim heights (Hr). B) Relative rim height (Hr/D).

Figure 4: The crater diameter above which central peaks are observed (s-c transition diameter) is inversely proportional to the planetary gravity. The general trends for the rocky and icy bodies are marked with separate lines. The transition diameter for icy bodies follows a similar trend, but at crater diameters an order of magnitude less than for rocky bodies. The icy body trend line is based only on the most recent high-resolution work using Galileo and Cassini images [Schenk, 1992; Croft et al., 1995; Schenk 2002, White et al., 2013]. Lower resolution results from Voyager (Schenk 1989) are shown with grey data points for comparison to demonstrate how changes in image resolution and coverage can affect the observed s-c transition diameter. Predicted minimum transition diameter for Pluto is marked with a black data point. The arrow extending from the Pluto data point marks that this prediction should be considered a lower bound.

Figure 5. Examples of impact-related pitting observed in craters: A) small (~ 1 km) terraced crater on the Moon at 302.8E, 8.3 N, B) a 22 km diameter summit-pit crater on Mars at 304.6E, 5.7N, C) Isis, a 70 km diameter floor-pit crater on Ganymede at 201.5E, 67.5S, D) a 3.5 x 3.5 km image of small-scale pitting on the floor of martian crater Corinto which is imaged in E and F. E) PSP_03611_1970 - a full-resolution (50cm/pixel) HiRISE image of Corinto crater on Mars at 141.6E, 16.9N. Small-scale pitting can be noted surrounding the main floor-pit. F) A 100m/pixel version of PSP_03611_1970.

Figure 6. Examples of high heat flow crater morphologies. A) Eshmun, a 96.5 km diameter crater on Ganymede with a central dome. B) Memphis Facula, a 355km diameter palimpsest on Ganymede. C) Valhalla, a multi-ring basin on Callisto with a 360km diameter central palimpsest and ring system reaching 1900km from the feature center. D) Tyre, a ~ 30km diameter multi-ring structure on Europa. This is significannlty smaller than expected for the occurrence of multi-ring basins and is the result of impact into a possibly thin, high heat flow icy crust. These crater morphologies are not expected to be observed by New Horizons unless Pluto’s crustal heat flow has been or is significantly higher than current predictions.

Highlights • • • •

Pristine impact crater morphology on Pluto is predicted Hydrocode modeling is used to study the effect of impact velocity Observations of craters on Pluto could help asses central pit formation Crater morphology can reveal details about the Pluto sub-surface