Europlanet Science Congress 2020
Virtual meeting
21 September – 9 October 2020
Europlanet Science Congress 2020
Virtual meeting
21 September – 9 October 2020

Oral presentations and abstracts

TP4

Impact processes shaped the solar system and modify planetary surfaces until today. This session aims at understanding planetary impact processes at all scales in terms of shock metamorphism, dynamical aspects, geochemical consequences, environmental effects and biotic response, and cratering chronology. Naturally, advancing our understanding of impact phenomena requires a multidisciplinary approach, which includes (but it is not limited to) observations of craters, strewn field or airbursts, numerical modelling, laboratory experiments, geologic and structural mapping, remote sensing, petrographic analysis of impact products, and isotopic and elemental geochemistry analysis.

We welcome presentations across this broad range of study and particularly encourage work that bridges the gap between the investigative methods employed in studying planetary impact processes at all scales.

Convener: Robert Luther | Co-conveners: Natalia Artemieva, Christopher Hamann, Isabel Herreros, Elena Martellato, Jens Ormö, Francisco Javier Rodriguez Tovar

Session assets

Session summary

Chairperson: Robert Luther, Elena Martellato, Jens Ormo, Christopher Hamann, Francisco Javier Rodríguez Tovar, Isabel Herreros
Giant impacts
EPSC2020-217ECP
Nicole Güldemeister, Lukas Manske, Miki Nakajima, and Kai Wünnemann

Planetary collisions play an important role in the compositional and thermal evolution of the planetary system. The Moon-forming impact event is thought to be Earth’s last giant collision event, marking the end of the main accretion phase of the Earth. This large event (re)set the conditions for the subsequent thermochemical evolution of both bodies, Earth and Moon. Large parts of proto-Earth are thought to melt as a consequence of the impact and the extent of melting affects the evolution of the Earth’s interior and atmosphere. It is critical to address the initial conditions of the proto-Earth and the volume and shape of a possible magma ocean after the impact to understand the Earth’s subsequent evolution. To address these questions, we consider two different impact scenarios and present a benchmark study by comparing two different numerical codes. Further, we investigate the effect of strength on melt production for the two different impact scenarios. We compared results using the shock physics code iSALE [1,2], an Eulerian code with a fixed grid in space and a Langrangian mesh-free, so-called smoothed particle hydrodynamics code, (SPH [3]).

Therefore, we consider two different impact scenarios that are described in Nakajima and Stevenson (2015); the canonical model (i) (e.g., [4]) and a fast-spinning Earth model (ii) [5]. The former (i) assumes a Mars-size impactor colliding with proto-Earth at an impact angle of 41° (90° corresponds to head-on collision) at a velocity corresponding to the escape velocity of Earth. The colliding bodies are differentiated into a mantle and core. The latter (ii) case is characterized by a smaller impactor colliding with a rapidly rotating Earth at a steeper impact angle of 73° and a much higher impact velocity of 20 km/s. In iSALE the impactor and target mantle consist of dunite, the core of iron; both materials are represented by an Analytical Equation of state (ANEOS [6]), the SPH code uses an MgSiO3 liquid and ANEOS for forsterite and iron. In iSALE we neglect the core of the impactor as well as the initial spin of Earth. The pre-impact temperature of the colliding bodies is 2000 K and near the solidus. Nakajima and Stevenson (2015) only consider hydrodynamic behavior of the colliding bodies. In the iSALE models a constitutive model accounting for strength is included. Thus, in addition to a comparison of the two different codes, we investigate the effect of strength on melt production. The simulations of the two codes are compared in terms of thermodynamic parameters (pressure evolution, Figure 1) as well as melt production (Figure 2) after the moon-forming impact.

In Figure 1, a comparison of pressure evolution of both impact scenarios and the two numerical approaches is shown. In general, a good agreement in terms of pressure ranges for the two codes is observed. In the fast-spinning model a large fraction of mantle undergoes an extensive expansion and becomes subject to low pressure before some part falls back towards the core where it gains high pressures (Nakajima and Stevenson 2015). This process is not as prominent in the standard model. In the iSALE simulation we do not observe the fall back of the material although the post-impact state of Earth looks similar in both simulations.

The SPH simulations show that the majority of mantle experiences melting (between 80 and 100% mantle melting) pointing to the existence of a global magma ocean with a base close to the core-mantle boundary in both impact scenarios (Nakajima and Stevenson 2015).

The melt production obtained by the iSALE simulation is at least for the standard impact scenario smaller. For the fast-spinning Earth model about 60% of mantle material are molten. Thus, for the standard model, we only observe partial mantle melting whereas for the fast-spinning model we can assume that a global magma ocean exists. However, we have to note that the iSALE simulations have not run as long as the SPH simulations yet and that with a longer run time melt production most likely increases. Neglecting an impactor core in iSALE will also underestimate melt production. The accretion of the core would increase melting.

Regarding the effect of strength, we observe a slight increase of melting when the colliding bodies exhibit some initial strength. However, the effect of the chosen impact scenario is much more significant as seen in Figure 2. The melt efficiency (melt volume normalized by the impactor volume) is over three times larger for the fast-spinning Earth model (2.66 corresponding to 60% of mantle volume) than for the standard model (0.73 corresponding to 20 % of mantle volume).

The presented benchmark study shows an overall good agreement of two numerical codes and indicates the significance of the chosen impact scenario with respect to melt production on Earth during the Moon-forming impact. However, the differences of the two numerical methods may also be caused by some differences in the setup (e.g. different EoS, spin of Earth, differentiated impactor).

In addition, in an ongoing study we systematically carry out simulations of the moon-forming impact event using iSALE. We vary the impact angle (15° to 60°) and impact velocities (12 to 20 km/s) and use different initial temperature profiles. First results show that with increasing impact angle as well as with increasing impact velocity, melt production increases. The effect of the pre-impact temperature becomes more prominent for increased impact angles.

How to cite: Güldemeister, N., Manske, L., Nakajima, M., and Wünnemann, K.: Numerical modelling of the thermal state of Earth after the Moon-forming impact event - A benchmark study, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-217, https://doi.org/10.5194/epsc2020-217, 2020.

EPSC2020-1033ECP
Yonatan Shimoni, Oded Aharonson, and Raluca Rufu

Introduction: A common outcome of a giant impact event is the formation of a circumplanetary debris disk, and in some cases, the capture of the surviving impactor, which results in a system composed of a primary, secondary, and a debris disk. The material in the debris disk may be accreted by the primary body (and secondary, if it exists), escape the system, or coalesce into larger clumps. A classic example of a system thought to originate by such an impact event is the Pluto-Charon binary system (Canup, 2005, 2011), in which all six bodies (including the other smaller satellites) lie approximately on the same plane and have nearly circular orbits. 

Such impacts are simulated using smoothed particle hydrodynamics (SPH) codes including self gravity. The equation of state (EOS) governs the relationship among the thermodynamic variables of the simulated material. The specific EOS used in the simulation may influence the final post-impact structure. In order to quantify the typical differences,  we compare two approaches, one simple analytic and one tabulated EOS. Tillotson (Tillotson, 1962) is a widely used analytic EOS and is computationally fast, but it lacks important details such as the treatment of phase changes. Sesame (Bennett et al., 1978) is a  commonly used tabulated EOS  and is more accurate, but is computationally slower and may be poorly sampled in the required thermodynamic phase-space. Here we show a set of SPH impact simulations that assume similar geometric and dynamic initial conditions but different  EOS. 

Methods: We performed ~100 SPH simulations using SWIFT code (Schaller et al., 2018), simulating Pluto-like impacts with 105-106 particles. The initial bodies are assumed to be differentiated, with target to impactor mass ratio of 1 or 7/3, impact angle, ξ, of 0, 30, 45, 60°. Impact velocity was chosen to be relatively small, 1-1.1 times the escape velocity. The impactor was either spin-less or rotating with a period of 5 or 10 hours. This parameter space was motivated by previous simulations (Canup, 2011) for the formation of the Pluto-Charon system. We compare results using the Tillotson and Sesame EOS. The simulations were stopped after 4 days, the typical time for the central body to relax to a stable spherical shape. We developed an algorithm to detect post-impact clumps. Two particles were considered in contact if their mutual distance was smaller than their combined smoothing lengths. The orbital elements of each clump (“satellitesimal”, defined as 100 particles in pairwise contact, equivalent to ~10-3Mpluto) were computed and studied.

Results: Disk systems were formed for impact angles >30°, in using both EOSs. Satellitesimals, when formed, showed different properties. Figure 1 panels (a,b) show the final snapshots of an impact that produced a debris disk and several satellitesimals highlighted in color corresponding to their masses. In this example, a more massive debris disk with a larger number of satellitesimals is obtained when using Tillotson EOS than using Sesame. Moreover, in addition to the target body, at least one large satellitesimal was formed in each of the simulations, but their composition, mass, and orbital elements differ between the EOSs. In comparing the Sesame run to Tillotson, the largest satellitesimal has a mass of 0.015Mpluto (with water fraction of 0.16) and 0.008Mpluto (with water fraction of 0.30) respectively. Its orbital elements are e=0.67, a=10.07Rpluto for Sesame runs, and e=0.15, a=3.2Rpluto for Tillotson. At smaller masses, using Tillotson EOS produces a greater number of clumps, as seen in Figure 1c. Three more satellitesimals were formed (some beyond the plot limits of 1b), with e=0.86, 1.35, 0.10. Note that the satellitesimal mass is  an order of magnitude  smaller than Charon, so alone they do not predict Charon’s formation (Canup, 2011). 

In the final snapshot, a greater fraction of debris disk particles lie within the Roche limit (computed using present-day Charon’s density (McKinnon et al., 2017)) in the Sesame simulation, whereas the Tillotson disk extends to a greater distance. In terms of mass, angular momentum, and composition, the debris disks are similar (0.025 versus 0.033 of the total mass; 0.19 versus 0.25 of the total angular momentum; water fraction of 0.44 versus 0.37 for Sesame and Tillotson respectively). We note the execution time was 2-3 times longer for the simulations using Sesame than Tillotson, a factor which may be considered in choosing EOS.

Exploring the parameter space, we note that head-on impacts (ξ~0°) produce a merged single body, with the vast majority of ejected particles accreted by the planet, and the rest ejected to space. Oblique, faster than escape velocity impacts (vimp/vesc=1.1 and ξ=60°) resulted in two unbound bodies, with little mass transfer between the two, consistent with previous studies of planetary impacts (Leinhardt & Stewart, 2012).

 

References:

Bennett, B. I., Johnson, J. D., Kerley, G. I., & Rood, G. T. (1978). Recent developments in the Sesame equation-of-state library. https://doi.org/10.2172/5150206

Canup, R. M. (2005). A Giant Impact Origin of Pluto-Charon. Science, 307(5709), 546–550.

Canup, R. M. (2011). On a Giant Impact Origin of Charon, Nix, and Hydra. The Astronomical Journal, 141(2), 35.

Leinhardt, Z. M., & Stewart, S. T. (2012). Collisions Between Gravity-Dominated Bodies. I. Outcome Regimes and Scaling Laws. The Astrophysical Journal, 745(1), 79.

McKinnon, W. B., Stern, S. A., Weaver, H. A., Nimmo, F., Bierson, C. J., Grundy, W. M., et al. (2017). Origin of the Pluto–Charon system: Constraints from the New Horizons flyby. Icarus. https://doi.org/10.1016/j.icarus.2016.11.019

Schaller, M., Gonnet, P., Chalk, A. B. G., & Draper, P. W. (2018, May 1). SWIFT: SPH With Inter-dependent Fine-grained Tasking. Astrophysics Source Code Library. Retrieved from https://ui.adsabs.harvard.edu/abs/2018ascl.soft05020S

Tillotson, J. H. (1962). Metallic equations of state for hypervelocity impact (No. Rep. GA-3216 ). General Dynamics San Diego CA.

 

How to cite: Shimoni, Y., Aharonson, O., and Rufu, R.: Properties of Debris Disk and Satellitesimals in Pluto-Like Impacts Under Different Equations of State, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-1033, https://doi.org/10.5194/epsc2020-1033, 2020.

EPSC2020-665ECP
Harry Ballantyne and Martin Jutzi
 

Introduction

Almost 50 years ago, NASA’s Mariner 9 space probe performed the first complete orbit of a planet other than Earth. This was around our smaller celestial neighbour – Mars. Along the way, this spacecraft obtained images of approximately 85% of the planet’s surface in unprecedented detail, revealing a stark contrast between the northern and southern hemispheres now known as the martian crustal dichotomy. This moniker predominantly refers to the 4-8 km difference in elevation between the southern hemisphere and an apparent basin covering roughly 42% of the north. Other associated features include a higher density of volcanoes and visible impact craters in the south relative to the north.

Some studies have attempted to explain these properties through endogenic means; namely via degree-1 mantle convection studied using large-scale thermochemical models (e.g. Keller and Tackley, 2009). Others have taken the exogenic route, proposing that a giant impact early in Mars’ history caused the excavation of a large mass of material from the northern hemisphere, thus giving rise to the observed dichotomy (e.g. Marinova et al., 2008). Given that such collisions are expected to be very common in the final stages of terrestrial planetary accretion, this approach is highly feasible. The latter studies have, however, generally ignored any long-term geodynamical consequences on the martian interior that such an event may cause.

Recent work has proved the importance of coupling these methods, introducing a hybrid exogenic-endogenic scenario whereby a giant impact triggered a localized magma ocean and subsequent superplume in the southern hemisphere (Golabek et al., 2018). This hypothesis has, however, only been investigated using a very limited range of initial parameters, all of which lead to significant heating deep into the mantle. This therefore motivates an interesting area of study – could there be a parameter space that leads to a hemispherically-thickened crust without significantly heating the mantle? We aim to answer this question using a suite of smoothed-particle hydrodynamics (SPH) simulations that explore a large parameter-space chosen with the intention of limited internal heating, allowing us to neglect any long-term geodynamical effects with reasonable confidence.

Method

The chosen initial parameters for the main suite of SPH simulations are as follows: impact angles of 0-90° in steps of 15°; impact velocities of 1.0, 1.2 and 1.4 times the mutual escape speed; impactor radii of 1000km, 1500km and 2000km; and relative core masses of 25% and 50%. All of these simulations use a resolution of 200,000 SPH particles, with near head-on collisions (0-30°) being modelled for 50 hours after impact and oblique collisions (45-90°) for 200 hours to allow for any secondary (or even tertiary) impacts. In addition, a smaller set of high resolution (1,000,000 SPH particle) simulations are being used to investigate extremely low-velocity (less than mutual escape speed) collisions in a bid to quantitively identify the transition from impacts of large-scale mantle heating to those of relatively cool accretionary piles similar to those described in Jutzi and Asphaug (2011). This then allows for discussion of the potential mechanisms that could lead to such events and their feasibility.

Each model includes the effects of shear strength and plasticity (via a Drucker-Prager-like yield criterion) as such effects have been shown to be significant on the scales concerned in this study (Emsenhuber et al., 2018). Moreover, the sophisticated equation of state ANEOS is being used along with a Mars-specific solidus (Duncan et al., 2018) to accurately calculate the physical environment in which such solid characteristics must be considered.

In all of these studies, both Mars and the impactor are treated as differentiated bodies composed of an iron core and a silicate mantle. An example result of one of these simulations can be seen in Figure 1.

Figure 1: The resulting temperature field of a Mars-like body from a 106 particle SPH simulation 12hr after an impact with a ≈1000km radius impactor at a 0° impact angle and a velocity of 80% of the mutual escape speed.

Preliminary Results

The initial results of this study have revealed promising hemispherical features in certain cases, particularly when examining the melt distribution in the upper mantle. Of notable interest are the results of the grazing impact angles, as some of the initial kinetic energy of the impactor is converted to rotational energy of Mars, allowing for subsequent merging events of decreased impact velocity. In addition, the effects of material strength have been found to be non-negligible, in contrast to previous beliefs that such aspects can be ignored on the length-scales involved in planetary collisions.

References

Canup, R. and Salmon, J. (2018). Origin of Phobos and Deimos by the impact of a Vesta-to-Ceres sized body with Mars. Science Advances, 4(4).

Duncan, M. S., Schmerr, N. C., Bertka, C. M., and Fei, Y. (2018). Extending the Solidus for a Model Iron-Rich Martian Mantle Composition to 25 GPa. Geophysical Research Letters, pages 211–220.

Emsenhuber, A., Jutzi, M., and Benz, W. (2018). SPH calculations of Mars-scale collisions: The role of the equation of state, material rheologies, and numerical effects. Icarus, 301:247–257.

Golabek, G. J., Emsenhuber, A., Jutzi, M., Asphaug, E., I., and Gerya, T. V. (2018). Coupling SPH and thermochemical models of planets: Methodology and example of a Mars-sized body. Icarus, 301:235–246.

Jutzi, M. and Asphaug, E. (2011). Forming the lunar farside highlands by accretion of a companion moon. Nature, 476(7358), 69-72.

Keller, T. and Tackley, P. J. (2009). Towards selfconsistent modeling of the martian dichotomy: The influence of one-ridge convection on crustal thickness distribution. Icarus, 202(2):429–443.

Marinova, M. M., Aharonson, O., and Asphaug, E. (2008). Mega-impact formation of the Mars hemispheric dichotomy. Nature, 453(7199):1216–1219.

Murchie, S. L., Thomas, P. C., Rivkin, A. S., and Chabot, N. L. (2015). Phobos and Deimos.

How to cite: Ballantyne, H. and Jutzi, M.: Forming the Martian Crustal Dichotomy Without Significant Mantle Heating, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-665, https://doi.org/10.5194/epsc2020-665, 2020.

Processes and products of impact cratering
EPSC2020-1095
Elliot Sefton-Nash, Olivier Witasse, and Zoe Faes

1. Introduction

Moonlets from an equatorial debris disk caused by a giant impact (a mechanism by which Phobos and Deimos may have formed (Craddock 2011; Rosenblatt 2011), may have had slowly decaying orbits leading to craters formed at low impact angles (< 5°). The absence of crater morphologies consistent with very low impact angles has been used to argue against the spiralling moonlet hypothesis for the formation of elongated craters (Bottke et al. 2000). The absence of comprehensive atmospheric entry and impact modelling for decaying moonlets leaves open the possibility that some only moderately elongated craters on Mars were formed by decaying moonlets in a thicker atmosphere. Indeed, the 12.5km spatial separation and apparent 3-3.7Ga age and cogenesis of double-oblique impact craters observed by (Chappelow and Herrick 2008) could be inconsistent with formation by a fast (i.e. non-moonlet) meteorite impactor unless a thicker atmosphere provided drag to increase impact angle during spiralling.

2. Method

To explore the effect of ancient atmospheric conditions on the impact angle of potential spiralling moonlets, we construct a simple 2D model to spiral down a moonlet from the top of a hypothetical Mars atmosphere. A body that is at mean circular orbital velocity at the top of the atmosphere,

experiences gravitational pull

and atmospheric drag,

Where G is the gravitational constant, Mm is the mass of Mars, r is the distance between the centres of mass of the impactor and Mars’, CD is the coefficient of drag, 𝜌 is the atmospheric density, and A is the cross-sectional area of the impactor (with respect to the velocity direction).

The model proceeds at each time step (1s) by updating the impactor velocity with a vector of the resultant acceleration, calculated using instantaneous gravitational and atmospheric drag forces. We simulate impact angle, retrieved at the point of contact between spherical impactors and the Mars surface for plausible parameter ranges for impactors and ancient Mars atmosphere.

CD is typically determined experimentally, and is a function of many parameters. At low speeds and for a small spherical projectiles, CD  ~ 0.5 (Miller and Bailey 1979). For large projectiles and mach numbers > 1, behaviour of CD becomes complex and may reach values > 1. To account for this we run simulations for a range of CD up to 1.1.

Impactors 10m to 10km in diameter, surface atmospheric pressures 5 to 500 mbar, coefficients of drag 0.5 to 1.1, and densities covering the range between Phobos and Deimos (1400-1900 kg m-3). An atmospheric temperature profile from  (Schofield et al. 1997) was used, and linearly interpolated as a function of altitude.

We assumed a present day atmospheric composition, with component fractions of 0.7532, 0.227, 0.016 and 0.08 of CO2, N2, Ar and CO, respectively. It is highly likely that an ancient Mars atmosphere had a different composition. As a cursory assessment of the potential impact of different atmospheric composition, we note that substituting 20% of Mars CO2 fraction for N2, the mean molecular mass, 𝜇, would change from 0.0455 to 0.0415, a shift of ~9%. Given that 𝜇 appears only a first order denominator in the calculation of scale height (H = RT/𝜇g) we do not expect uncertainty on past atmospheric composition to have a major impact on our conclusions.

A summary of the parameters used is shown in Table 1. We ran the model for all permutations of the parameter grid, yielding ~2.6x105, model runs.

Table 1: Domain and sampling intervals for parameter space explored for spiralling impactors.

Figure 1: Impact angle as a function of surface atmospheric pressure (mbar) and ballistic coefficient (BC). Surface is a bilinear interpolation of BC run over the parameters listed in Table 1.

3. Discussion

With a drag coefficient of 0.5, Phobos has a ballistic coefficient ~ 8x107 kg m-2, while Deimos’ is ~2.4 x 106 kg m-2. Both moons are therefore sufficiently large and dense that their impact angle would not be substantially modified by atmospheric drag, even for a 0.5 bar atmosphere. However, moonlets decaying from a quasi-stable equatorial debris disk has been proposed to be in the range 1-4km (Rosenblatt et al. 2016). A 1km diameter object with Deimos density would have a BC ~1.8x106. Such an object would experience a shift in impact angle of several 10s of degrees in atmospheres > 100 mbar.

Importantly, we note that the model does not account for non-spherical shapes, heterogeneous density distribution in the impactor, or fragmentation of impactors. Fragmentation would produce a group of impactors, each with lower ballistic coefficients than the parent impactor, and therefore whose impact angle may be steeper than if fragmentation did not occur.

4. Conclusions

We find that apparent lack of very oblique impacts that should be expected from decaying moonlet hypothesis could be due to increased atmospheric drag in an ancient, thicker atmosphere. Based on a simplistic model, impact angle deviations > 10° (compared to present day) are possible in atmospheres of few 100 mbars, for objects with low ballistic coefficients in the ~sub-km diameter range.

5. References

Bottke, W. F., Love, S. G., Tytell, D., and  Glotch, T. 2000. “Interpreting the Elliptical Crater Populations on Mars, Venus, and the Moon.” Icarus 145(1): 108–21.

Chappelow, J. E., and R. R. Herrick. 2008. “On the Origin of a Double, Oblique Impact on Mars.” Icarus 197(2): 452–57.

Craddock, R. A. 2011. “Are Phobos and Deimos the Result of a Giant Impact?” Icarus 211(2): 1150–61.

Miller, D. G., and Bailey, A. B. 1979. “Sphere Drag at Mach Numbers from 0·3 to 2·0 at Reynolds Numbers Approaching 107.” Journal of Fluid Mechanics 93(3): 449–64.

Rosenblatt, P. 2011. “The Origin of the Martian Moons Revisited.” Astronomy and Astrophysics Review 19(1).

Rosenblatt, P. et al., 2016. “Accretion of Phobos and Deimos in an Extended Debris Disc Stirred by Transient Moons.” Nature Geoscience 9(8): 581–83.

Schofield, J. T. et al. 1997. “The Mars Pathfinder Atmospheric Structure Investigation/Meteorology (ASI/MET) Experiment.” Science 278(5344): 1752–58.

How to cite: Sefton-Nash, E., Witasse, O., and Faes, Z.: Impact angle of impactors decaying from circular orbits in an ancient Mars atmosphere, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-1095, https://doi.org/10.5194/epsc2020-1095, 2020.

EPSC2020-708ECP
Andrea Rajsic, Katarina Miljković, Gareth Collins, Kai Wünnemann, Mark Wieczorek, Natalia Wojcicka, and Ingrid Daubar

Introduction: Meteoroid bombardment is one of the sources of seismic activity on planetary bodies. The very first seismometer operating on the surface of another planet was successfully deployed by the NASA InSight (Interior Exploration using Seismic Investigations, Geodesy and Heat Transport) mission to Mars. It gives us an opportunity to investigate the seismicity of Mars, including impact-induced seismic activity. This work investigated the seismic efficiency associated with small meteorite impacts on Mars, using numerical methods in targets analogue to the Martian surface. The Martian crust was simulated as non-porous bedrock (0% porosity) or regolith with different porosities (25%, 44% and 65%)

The seismic efficiency, k, is presented as a portion of impact energy that is transferred into seismic energy. It has been suspected that consolidated (bedrock) and non-consolidated (regolith) materials will have different values of seismic efficiency. Estimates of seismic efficiency range from k=10-2 to 10-6 (Schultz and Gault, 1975; Daubar et al., 2018; McGarr et al., 1969; Hoerth et al., 2014; Richardson & Kedar, 2013; Güldemeister & Wünnemann, 2017). High seismic efficiency is typical in bedrock or highly consolidated materials (k>10-3). Low seismic efficiency is typical for sediments or unconsolidated sands and soils (k<10-5) (e.g., Patton and Walter 1993). In this work, we used a simplified approach (e.g. Güldemeister & Wünnemann, 2017) that defines the seismic efficiency as: ; where x represents distance from the impact point, P is the amplitude of the pressure pulse, t is the duration of the pressure pulse, ρ is the density of the target, Cp is the speed of sound in the target and Ek is the kinetic energy of the impactor.

Numerical impact modelling: All simulations were performed with the iSALE-2D shock physics hydrocode (Collins, et al., 2004; Wünnemann et al., 2006). The impact conditions were modelled to replicate recent fresh meter size impact that occurred on Mars since the landing of InSight (Daubar et al., 2020). Impact crater was estimated to be ~1.5 m in diameter. Impactor radius was 4.4 cm and kinetic energy of 1.8x106 J.

To simulate bedrock and fractured bedrock (25% porosity) we used the ROCK strength model (Collins et al., 2004). To simulate the regolith (44% and 65% porosity) we used Lundborg strength model (Lundborg, 1968) (Table 1). We used the Tillotsen equation of state for basalt (Tillotson, 1962; Wójcicka et al., 2020). For porous cases, we used the ε-α porosity model (Wünnemann et al., 2006;) (Table 2).

 

Table 1. Strength model parameters for targets with different porosity

Parameter 0% 25% 44%  65% 
Strength model ROCK ROCK LUNDD LUNDD
Strength (damaged) (kPa) 10 0 10 0.3
Friction (damaged) 0.6 0.67 0.7 0.7
Limiting strength (damaged) (GPa) 3.5 0.17 0.25 0.25
Strength (intact) (MPa) 10 0.2    
Friction (intact) 1.2 1.8    
Limiting strength (intact) (GPa) 3.5 0.17    

 

Table 2. ε-α porosity model parameters (Borg et al., 2005; Wünnemann et al., 2006; Wójcicka et al, 2020).

Parameter 25% 44% 65%
Initial distension, α 1.33 1.8 2.8
Elastic threshold, ε0 -4x10-4 10-4 10-5
Distension at transition αx 1.1 1.15 1.0
The rate change of distension with respect to volumetric strain, k 0.98 0.98 0.98
Ratio of speed of sound in porous over non-porous medium, χ 0.6 0.33 0.21

 

All variables in the equation for the seismic efficiency were calculated from iSALE outputs. The pressure wave was observed via gauges cells, placed at 45° equidistantly throughout the target. The pressure wave amplitude and pulse duration were calculated at full width half maximum. The sound speed was calculated from assumed bulk modulus of basalt (Wójcicka et al, 2020).

Results: Seismic efficiency was calculated for the same impact conditions in all four material models, representing the reference 1.5 m crater recently observed on Mars. There is a clear decrease in seismic efficiency with increasing porosity. It is of the order of 10-5 for porous and highly porous regolith and 10-4 for fractured bedrock. Estimates for the non-porous basalt bedrock are in the order of  10-3 (Figure 1).

Porosity of the target affected pressure wave amplitudes, duration of the pressure pulse and speed of sound in the target. These are all parameters used in calculation of seismic efficiency. This implies that if impact occurs on very dusty parts of Mars with thick regolith cover, efficiency would be smaller than efficiency of the impact that occurred on the bedrock, or area with thinner regolith cover.

Figure 1. Seismic efficiency calculated in targets with different porosity

Conclusions. Impact cratering represents one of the most important geological processes in the Solar System. Defining relationship between target’s properties and seismic efficiency is of interest to the NASA InSight science, since it helps in understanding the properties of the uppermost crust on Mars. Previous approximations of seismic efficiency were of 2x10-5 with an order of magnitude uncertainty for seismic efficiency on Mars (Teanby and Wookey, 2011) and Daubar et al. (2018) adopted the seismic efficiency of 5x10-4 calculated from the seismic moment (Gudkova et al., 2011; 2015; Teanby 2015). In this work, we calculated the seismic efficiency in meter-size impacts on Mars to be different for bedrock (order of 10-3) and porous materials (order of 10-5).

How to cite: Rajsic, A., Miljković, K., Collins, G., Wünnemann, K., Wieczorek, M., Wojcicka, N., and Daubar, I.: Seismic efficiency of Martian upper crust simulant., Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-708, https://doi.org/10.5194/epsc2020-708, 2020.

EPSC2020-869ECP
Christopher Hamann and Lutz Hecht

Introduction

Understanding the shock behavior of the calcium sulfates gypsum (CaSO4·2H2O), bassanite (CaSO4·0.5H2O), and anhydrite (CaSO4) is essential for understanding hypervelocity impacts into evaporite sediments that are widespread on Earth and which also occur on Mars [1,2]. Most interest focuses on quantification of the impact-induced release of volatiles (H2O and SO2/SO3) to assess its role in the generation or modification of planetary atmospheres [3–11]. Estimates of the amount of gas released from volatile-bearing materials in hypervelocity impact scenarios rely on assignment of the shock-pressure thresholds for incipient and complete vaporization [10]. However, these thresholds are poorly constrained for calcium sulfates, and many studies produced inconsistent to even contrasting results (cf. [5,7,9]). In addition, reports of sulfate impact melts [6] suggest that the shock behavior of the calcium sulfates is more complex than traditionally thought and possibly involves interaction with coexisting silicate impact melts.

Based on our previous study on the interaction between carbonates and silicates in laser-generated melts [12], we report here on the fate of gypsum and anhydrite in laser-irradiation experiments. Specifically, we laser-irradiated gypsum, gypsum–granite, and anhydrite–granite targets to investigate dehydration, decomposition, and melting of the calcium sulfates as well as interaction between calcium sulfates and silicates in an impact-related context.

Materials and methods

We used a continuous-wave fiber laser of 1.07 µm wavelength at Fraunhofer Institut für Kurzzeitdynamik, Freiburg, Germany, to irradiate a gypsum plate (experiments L1130 and L1131) and pulverized gypsum–granite (experiment L1133) and anhydrite–granite (experiment L1132) mixtures in air at 1 bar and room temperature (Fig. 1). The laser emitted a power of 2 or 5 kW for 2 or 5 s, and 1/e2 beam diameters of 6 or 17 mm were used. The laser intensity varied between 8.8 × 102 and 1.8 × 104 W/cm2 and the laser–matter interaction zone was observed using the setup described in [13]. Temperature measurements (Fig. 2) show that peak temperatures were in the range of 1200 K for the gypsum plate and 2200 K for the composite targets. Textural and compositional characterization of the recovered samples employed transmitted-light microscopy, Raman spectroscopy, and scanning electron microscopy.

Results and discussion

Laser irradiation of the gypsum plate resulted in dehydration of gypsum and successive transformation of gypsum to anhydrite via bassanite. The transformation is evident from the 3250 to 3750 cm–1 region in Raman spectra (Fig. 3), in which the Raman bands corresponding to the stretch vibration modes of water are positioned. Consistent with previous investigations [8,14], the gypsum to anhydrite transformation is also accompanied by a slight upshift of the Raman band associated with the ν1 symmetric stretch vibration modes of the SO4 tetrahedra (i.e., from 1010 cm–1 in gypsum to 1017 cm–1 in bassanite to 1018 cm–1 in anhydrite).

In addition, dehydration of gypsum is also evident from BSE images obtained from the irradiation zones by an increase in brightness (Fig. 4a) and by the presence of cracks in individual anhydrite crystals that are absent in the gypsum crystals of the starting material (Figs. 4b,c). In addition, textures indicative of melting (occurrence of round vesicles; filling of cleavage space by flow-textured material of anhydrite composition) are present in the interior parts of the laser-irradiated regions (Fig. 4d).

Textures indicative of melting and recrystallization of anhydrite were also detected in clast-bearing melt droplets recovered from the gypsum–granite and anhydrite–granite experiments (Fig. 5). Specifically, anhydrite intimately associated with silicate melts occurs in the form of clusters of lath-shaped, typically vesicle-bearing crystallites that often enclose small silicate melt droplets. These clusters are interpreted as quenched CaSO4 liquids. The interfaces between the silicate melts and the anhydrite clusters are typically sharp and reminiscent of menisci formed by liquid immiscibility (cf. [8]; Fig. 5).

Furthermore, chemical interaction between CaSO4 and the silicate melts derived from the granite is evident from elemental distribution maps (Fig. 5) and point analyses (Fig. 6). Specifically, most silicate melts are heavily enriched in CaO (Fig. 6) as well as SO3, showing data trends in binary (e.g., CaO–SiO2) or ternary (e.g., CaO–SiO2–Al2O3) composition space that cannot be explained by melting and mixing of the granite components alone. In keeping with previous findings [6], limited amounts of SiO2 were detected in the quench anhydrite, suggesting that the CaSO4 liquids and silicate liquids readily exchanged chemical components during the brief time scales (seconds; cf. Fig. 2b) in which the materials were in the liquid state.

Conclusions

Our results suggest that the thermally triggered dehydration of gypsum due to hypervelocity impacts or magmatic activities can be exceptionally fast (seconds to tens of seconds) and results in dehydration of gypsum, formation of anhydrite via basanite, and melting of anhydrite. The presented Raman spectra and microtextures may be useful for interpretation of similar reaction products observed on Mars [2,15–17]. Our results confirm previous observations on shock-melted sulfates [6] and suggest that impact melting of calcium sulfates should be indeed more prevalent than previously thought. Specifically, our experiments illustrate that calcium sulfate impact melts will readily interact with coexisting silicate impact melts, resulting in Ca-enriched silicate melts and Si-enriched, quenched anhydrite melts.

References

[1] Langevin et al. 2005, Science 307, 1584–1586. [2] Vanimann et al. 2018, Am. Mineral. 103, 1011–1020. [3] Chen et al. 1994, Earth Planet. Sci. Lett. 128, 615–628. [4] Pierazzo et al. 1998, J. Geophys. Res. 103, 28607–28625. [5] Gupta et al. 2001, Earth Planet. Sci. Lett. 188, 399–412. [6] Osinski and Spray 2003, Earth Planet. Sci. Lett. 215, 357–370. [7] Skála et al. 2005, GSA Special Paper 384, 413–425. [8] Bell and Zolensky 2011, 42nd LPSC, Abstr. #2008. [9] Prescher et al. 2011, Meteorit. Planet. Sci. 46, 1619–1629. [10] Artemieva et al. 2017, Geophys. Res. Lett. 44, 10180–10188. [11] Kurosawa et al. 2019, Geophys. Res. Lett. 46, 7258–7267. [12] Hamann et al. 2018, Meteorit. Planet. Sci. 53, 1644–1686. [13] Hamann et al. 2016, Geophys. Res. Lett. 43, 10602–10610. [14] Liu et al. 2009, 40th LPSC, Abstr. #2128. [15] Knauth et al. 2005, Nature 438, 1123–1128. [16] Squyres et al. 2012, Science 336, 570–576. [17] Robertson and Bish 2013, Icarus 223, 407–417. 

How to cite: Hamann, C. and Hecht, L.: Dehydration, melting, and recrystallization of calcium sulfates (gypsum, bassanite, anhydrite) in laser-irradiation experiments, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-869, https://doi.org/10.5194/epsc2020-869, 2020.

EPSC2020-1027
Gabriele Giuli, Sabrina Nazzareni, and Henrik Skogby

The Dellen impact structure is located in a ca. 20 km wide basin filled by two lakes in the east central Sweden. This complex structure is covered by a thick moraine deposit, and most impact melted material can be found as loose blocks and boulders scattered throughout the moraine. Geophysical measurements show the presence of a coherent impact melt body about 9 km in diameter and 200 to 500 m thick (Henkel, 1992) for which impact ages ranging between 90, 110, to 140 Ma have been determined (see Deutsch et al., 1992; Mark et al., 2014, and references therein).

Here we describe the petrography and mineral chemistry of an impact melt glass (commonly referred to as dellenite) consisting of phenocrysts of subhedral orthopyroxene (average Wo4En63Fs33), skeletal plagioclase (average An59Ab38Or3), and euhedral magnetite within a glassy matrix of rhyolitic composition.

Moreover we aimed to use the pyroxene crystals as a petrological tool to define the host rock cooling rate and the water content dissolved in their parental melt during pyroxene growth.

MicroFTIR measurements performed both in the glass matrix and in the euhedral glass blebs within the skeletal plagioclase show the presence of 1.4 wt% water in the rhyolitic glass. Comparing the water content of the glass with the water solubility of a silicate melt with the same composition as determined by the model by Papale, suggests that the studied sample vitrified at a pressure of about 200-300 bar.

Orthopyroxene crystals from the dellenite sample were studied by polarized FTIR spectroscopy, Mössbauer spectroscopy, single crystal X-ray diffraction and electron probe micro analysis. Dellenite orthopyroxenes are iron rich enstatite (Wo4 En63 Fs33) with a Fe3+/Fetot ratio of 1.6%, as measured by Mößbauer spectroscopy.

The studied orthopyroxenes have very weak to absent OH vibrational bands in the IR spectra, corresponding to H2O contents ranging from 0 to 39 ppm H2O. These variable contents seem to indicate H loss during post-formation processes, which may occur via the relatively fast redox reaction Fe2+ + OH- = Fe3+ + O2- + ½H2.

In volcanic pyroxenes H loss may be a common process occurring at different moments from crystallization to post-eruption. However, H incorporation in pyroxene is associated with point defects governed by slow kinetics diffusion rates, which are retained in the structure when H is lost. By reversing the redox reaction the original H content of pyroxene can thus be recovered by thermal annealing experiments under reducing conditions (Weis et al., 2016).

In order to restore H that was possibly lost, dellenite pyroxenes were thermally annealed under hydrogen atmosphere (at 1 Atm) in a horizontal glass-tube furnace at the Department of Geosciences (Natural History Museum Stockholm). We performed thermal annealing experiments at 700°C for 17 hours. FTIR spectra were recorded after each heating step. All the samples increased their hydrogen content and the final average water content is 77 ppm.

Cation partition as derived by the single crystal X-ray diffraction and EMPA were used to calculate the orthopyroxene closure temperature (Tc) which is expression of the cooling rate for the cpx-host rock. Preliminary results point to a high Tc (833°C) due to a quite fast host-rock cooling rate.

A tentative model for the evolution of the rhyolitic portion of the Dellen impact melt is under development merging geophysical data with our experimental data on water content of pyroxene and glass, and pyroxene geospeedometry.

 

 

References:

- Deutsch A., Buhl D., Langenhorst F. (1992) on the significance of crater ages: new ages for Dellen (Sweden) and Araguainha (Brazil). Tectonophysics, 216, 205-218.

- Henkel H., (1992) Geophysical aspects of meteorite impact craters in eroded shield environment, with special emphasys on electric resistivity. Tectonophysics, 216, 31-40.

- Mark D.F., Lindgren P., Fallick A.E. (2014) A high precision 40Ar/39Ar age for hydrated impact glass from the Dellen impact, Sweden. Geological Society of London Special Publication 378, 349-366.

- Papale P., (1999) Modelling the solubility of H2O+CO2 fluid in silicate liquids. American Mineralogist, 84, 477-792.

- Weis FA, Skogby H, Stalder R (2016) Experimental hydration of natural volcanic clinopyroxene phenocrysts under hydrothermal pressures (0.5–3 kbar). American Mineralogist 101(10): 2233-2247

How to cite: Giuli, G., Nazzareni, S., and Skogby, H.: Dellen impact melt rock: pyroxenes as a proxy for the melt thermal history and water content, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-1027, https://doi.org/10.5194/epsc2020-1027, 2020.

EPSC2020-235
Agnese Fazio and Luigi Folco

1. Introduction

A recent study on a Muong Nong-type tektite [1] revealed the occurrence of quartz with a “spongy texture”. The orientation and distribution of the voids in quartz were supposed to be crystallographically controlled.

Similar structures were not reported before in the literature except for a glass impactite from the 45-m-Kamil crater, Egypt (sample L09) [2]. No Raman or TEM studies were carried out on this lapillus and due to the occurrence of voids, this SiO2-rich material was interpreted as glass.

The renewed interest induced us to check again these structures in Kamil impactites. They were found along borders and fractures in numerous quartz relicts occurring in impact melt lapilli. In this work, we present the preliminary results of the characterization of these structures.

2. Impactite L09

The present study focuses on the SiO2 glass-rich impactite L09 (max length ~5 cm). This sample was described in [2] as white glass, i.e., a glass derived from the melting of the target rock without interaction with the projectile (Ni below detection limit). The glass is locally stained by reddish-brownish material. Among Kamil white glasses, L09 is the only sample showing large quartz relicts (up to a few millimeters).

3. Methods

The sample was optically characterized through polarized microscope and Raman spectroscopy. Successively, it was investigated by scanning electron microscopy (SEM) and spot chemical analyses from the vesicular glass and quartz relicts were acquired through electron microprobe (EMP).

4. Results

Under crossed polarized light the quartz relicts have a very low birefringence and show sporadic birefractive domains. In the birefractive domains, planar deformation features (PDFs) have been observed. Raman spectra of quartz relicts vary between a pure SiO2 glass spectrum with the typical large asymmetric band at ~490 cm-1 and a glass-free shocked-quartz spectrum (main peak at 460 cm-1). Shocked-quartz is easily recognized by Raman spectroscopy: the peaks are shifted towards lower frequencies due to the increase in the angle spanning two tetrahedra [3,4]. Up to two sets of PDFs were recognized at the SEM. PDFs are irregular, closely spaced, enlarged, and tend to coalesce. The PDF areas are in spatial continuity with PDF-free amorphous regions. The PDF-free amorphous regions can contain sporadic vesicles. No flux textures have been recognized in quartz relicts. Occasionally, quartz relicts are crossed by vesicular glass veins chemically enriched in Al2O3 and FeO and open fractures.

The margins of the quartz relicts and fractures are mostly surrounded by a layer (usually 30-µm-thick; exceptionally up to 100-µm-thick) of unshocked quartz (Raman main peak at 465 cm-1). Quartz appears in form of skeletal hexagonal aggregates. The numerous voids left out from the skeletal growth are partially filled by iron oxides. The number of the voids increases and their size decreases towards the quartz relicts describing a kind of rim. Small concentric fractures are frequent in the quartz relicts, marking the rim of the skeletal quartz layers. No cristobalite or high-pressure SiO2 polymorphs (i.e., coesite and stishovite) have been detected by Raman.

5. Preliminary discussion

The occurrence of skeletal quartz is indicative of a rapid crystallization of the quartz from a fluid phase. Due to the size of the impact event and size of the lapillus, a hydrothermal post-shock alteration could be ruled out. Thus, it is plausible that the quartz crystallized from a SiO2-rich melt. Due to its viscosity, this melt should have been formed at the same location where the skeletal quartz occurs. The melt formed at the expense of the highly-shocked-quartz/diaplectic-glass grains with a minor involvement of the Si-Fe-Al surrounding melt/glass. The melting should have occurred immediately after the pressure release, in the early decompression stage. The subsequent drop of the temperature induced the fast (skeletal) crystallization of the quartz. That probably started in the stability field of β-quartz (4.5 GPa<P<0.6 GPs). The crystallization propagated radially towards the glass from the grain boundaries and fractures. Changes in the rim texture of the skeletal quartz could be related to changes in growth rate. The iron-rich oxides in the cavities of the skeletal quartz represent the immiscible/residual melt of the quartz crystallization. The forthcoming focus ion beam milling (FIB) and a transmission electron microscopy (TEM) investigations aim to establish the crystallographic and phase relations between the quartz relicts and the skeleton quartz aggregates to constrain their formation and meaning in the framework of the impact melting process.

Acknowledgements

This work is supported by the Deutsche Forschungsgemeinschaft (DFG; FA 1599/1-1 to AF). Prof. Langenhorst is thanked for the access to the SEM/FIB and TEM facilities at the University of Jena founded via the Gottfried Wilhelm Leibniz prize (LA830/14-1). Dr. Kiefer (University of Jena) is acknowledged for technical assistance during the EMP measurements. The studied sample was collected during the 2010 geophysical expedition carried out within the framework of the 2009 Italian- Egyptian Year of Science and Technology and supported by the Italian Ministero degli Affari Esteri e Cooperazione Internazionale (MAECI)—Progetti di Grande Rilevanza.

References

[1] Glass, B.P., et al. (2020) Coesite in a Muong Nong-type tektite from Muong Phin, Laos: Description, formation, and survival. Meteoritics & Planetary Science. 55:253–273.

[2] Fazio, A., et al. (2016) Target-projectile interaction during impact melting at Kamil Crater, Egypt. Geochimica et Cosmochimica Acta 180:33-50.

[3] Fritz, J., et al. (2011) Shock experiments on quartz targets pre-cooled to 77 K. International Journal of Impact Engineering 38:440-445.

[4] Mcmillan, P.F., G.H. Wolf, and P. Lambert (1992) A Raman-Spectroscopic study of shocked single crystalline quartz. Physics and Chemistry of Minerals 19:71-79.

How to cite: Fazio, A. and Folco, L.: Skeletal crystallization of quartz from SiO2 impact melt, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-235, https://doi.org/10.5194/epsc2020-235, 2020.

EPSC2020-928
Sean Gulick, Gail Christeson, Naoma McCall, Joanna Morgan, and Jens Ormö

Introduction:   Orbital images of impact crater floors show heterogeneity in texture, relief of central structures, peak rings or terrace zones, and presence or absence of blocks. The absence of the 3rd dimension renders challenges in interpreting the thickness of impactites that fill these craters. Whereas texture in orbital images give clues to emplacement process, terrestrial craters benefit from seismic imaging and scientific drilling.

The Cretaceous-Paleogene (K-Pg, 66 Ma), 200 km Chicxulub impact crater, México, provides the unique opportunity to study crater formation processes related to a large impact [1]. Chicxulub is a multi-ring basin with an intact melt sheet, peak ring, and crater floor (Fig. 1) [1,2]. It is well preserved due to its youth and burial beneath 100s of meters of Cenozoic carbonates [1,2].

Figure 1.  Northwest orientied, time migrated seismic line (Line B), which was depth converted using 3D seismic refraction velocity model, crosses from near the crater center to inner ring faults of Chicxulub impact structure showing features with 3x vertical exaggeration. Line shows that impactites represent substantial thicknesses of material burying slump blocks within annular trough, capping the peak ring, and overlying the central melt sheet. Top K-Pg represents top of suevite layer drilled in IODP-ICDP Expedition 364 (see Figure 2). Figure modified from [2].

The 14.9 Ma, 25 km Ries impact crater, Germany, appears to represent a transitional crater form between central peak and peak ring crater with a collapsed central uplift. It has a well-preserved crater floor within the central crater due to burial by lacustrine sediments, but no clear indication of a melt sheet [3].

Insights from Scientific Drilling and New Geophysical Analyses:  The International Ocean Discovery Program with co-funding from the International Continental Scientific Drilling Project drilled into the Chicxulub peak ring in 2016 [4]. Hole M0077A (Fig. 2) recovered the uplifted crystalline rocks of the peak ring and the layered deposit that overlies it consisting of impact melt rock overlain by clast-rich impact melt rock and breccia, suevite, and finally sorted suevite with decreasing clast size upsection.  These impactites all have a low velocity (~3000 m/s or less) and high porosity (20-40%). Interpretation of this low-velocity sequence is that the lower portion was rapidly deposited onto and perhaps intermixed with impact melt rocks during impactoclastic flows and melt-water interactions. The upper sorted suevite is a resurge deposit in this marine target setting [5].

Figure 2.  Seismic reflection Line ChicxR3 within the Chicxulub crater overlain by full waveform inversion generated velocity model. The Chicxulub impact crater peak ring is overlain by a low-velocity zone that can be mapped into the adjacent annular trough which maps to a sorted suevite within IODP-ICDP Exp. 364 Hole M0077A core. White line representing depth reach during drilling (1335 meters below seafloor) where this low velocity layer represented ~90 m thick and was underlain by additional 40 m of more melt rich impactites. Figure modified from [5].

2D, full-waveform tomographic velocity images from the grid of seismic data recorded on a 6-km streamer across the Chicxulub impact crater allowing mapping of these units (Fig. 2). These images show both the lower and upper portions of the low-velocity layer present at Site M0077 thus allowing us to potentially map the thickness of both the resurge deposit and the lower more rapidly deposited layer. Prominent high-velocity zones are also visible that are interpreted to represent intact melt rock (Fig. 2) [2].

In the Ries Crater new high-resolution images highlight layered sequences within the central crater. One interpretation of  this package is that it may represent so-called crater suevite penetrated by the Nordlingen 1973 drill core, that we suggest were deposited in impactoclastic flows. In this drill core, suevite is overlain by a thin section of sorted suevite perhaps similar to the upper low velocity zone at Chicxulub. Within the annular trough, the suevite lays unconformably over bunte breccia which was transported inward during crater modification [3].

Impactites and Crater Floor: The Chicxulub crater floor is largely made of the sorted suevite resurge deposit expressed as the upper low velocity zone, which only occasionally is penetrated by a cropping out of deeper impactites. This may reflect the volume of ejecta and impactites formed within the 100 km diameter transient cravity at Chicxulub, that were then uniformly distributed within the final crater due to depth of resurge. The lower more rapidly deposited layer however varies with the greatest thicknesses in the central basin at Chicxulub which may provide insights for Ries.

The layered crater suevite at Ries may be equivalent to this lower unit of melt rich impactites wherein impactoclastic flows incorporate the melt potentially providing an explanation for the lack of an intact melt sheet in this smaller crater. In turn the 2 km thick impactites in the annular trough at Chicxulub (Fig. 1) may represent both ejecta curtain material similar to the bunte breccia at Ries and outward directed melt rock and suevite such that the suevite overlies the ejecta similar to at Ries.  The Ries crater floor may have been more heterogeneous with some areas exhibiting blocks of bunte breccia emergent through the suevite and the sorted suevite perhaps not being present throughout the crater. 

References: [1] Morgan et al. (2007) Nature, 390, 472-476 [2] Gulick et al. (2013) Rev Geophysics, 51, 31-52 [3] Stöffler et al. (2013) Meteoritics & Planetary Science, 48, 515-589, [4] Morgan J. V. et al. (2016) Science, 354, 878–882 [5] Gulick et al. (2019) Proc Nat Acad Sci.,116, 19342-19351.

How to cite: Gulick, S., Christeson, G., McCall, N., Morgan, J., and Ormö, J.: Impactite stratigraphy and depositional processes in the Chicxulub and Ries impact structures: What is a crater floor? , Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-928, https://doi.org/10.5194/epsc2020-928, 2020.

Impacts and planetary surface evolution
EPSC2020-280ECP
Daniel Wahl, Mark Wieczorek, Kai Wünnemann, and Jürgen Oberst

The upper portion of the lunar highland crust is known to show a wide range of porosities between 3 and 23% (Wieczorek et al., 2013). Impact cratering seems to be the primary mechanism for the porosity structure of the lunar crust as it can be observed today. 

As demonstrated in previous work by Soderblom et al. (2015), impact cratering changes the initial porosity structure of the upper crust. Considering the Bouguer gravity signal of lunar complex craters, it was shown that preimpact porosity of the target material controls the magnitude of the postimpact Bouguer gravity anomaly. Impacts in targets with initially low porosity were found to result in a negative Bouguer anomaly (indicating a higher porosity after impact). In contrast, impacts in target materials with high porosity result in positive Bouguer anomaly (indicating a lower final porosity). To avoid the contribution of the gravity signal caused by the uplift of mantle material, their study concentrates on complex craters with diameters smaller than 200 km.

We compute a global map of crustal porosity in high spatial resolution (analysis region of 3° and grid size of 0.75°), using most recent Gravity Recovery and Interior Laboratory (GRAIL) mission gravity field model, developed to degree and order 1500 (Park et al., 2015). By filtering the long wavelength portion of the measured gravity signal, our results are independent from bending of the crust-mantle boundary, allowing us to investigate the porosity of impact basins with diameters larger than 200 km.

In a first step, bulk density was estimated by analyzing the correlation between Bouguer gravity and topography. ­­­Using only the short wavelength portion of the measured gravity field signal and applying a Bouguer correction (using the correct bulk density), the result should be uncorrelated with the surface topography. Testing different densities for calculating the Bouguer correction and minimizing the correlation with topography, the bulk density of each small analysis region is determined. Relating the bulk densities with independently obtained grain densities, lateral variations in porosity of the upper crust are determined (Figure 1). Since the lunar maria possess a complex interior structure, with dense basalts (originating most likely from the upper mantle) overlying the anorthosite crust, we omit locations of basaltic maria in our consideration.

Overall, we studied the porosity of 40 individual impact basins located in the lunar highlands.  Specifically, we compute azimuthally averaged radial porosity profiles. For many basins, we found distinct porosity signatures, with regions of low porosity (i.e., high compaction) in their center and high porosity about one crater radius from the center. Larger basins are found to show a stronger pronounced porosity signature than smaller basins. Higher porosities could be explained by a more extensive fracturing of the target rock, caused by a stronger shock wave and/or a thicker blanket of excavated material during crater formation. Compact material in basin centers could be associated with larger quantities and depths of impact melt. Some basins show a less distinctive and muted porosity pattern, which we attribute to postimpact modification by superposed large basins and craters (Figure 2).  Moreover, even though the data statistics are limited, we tentatively identify a correlation between the porosity structure of impact basins and their relative ages. While older candidates show a less pronounced porosity signature with a lower porosity contrast, younger basins show the highest porosities near their rims and the lowest values interior to the peak ring. We relate the lower contrast for older basins to modifications through later impact events. We suggest that after formation all basins show a pronounced porosity signature, which is erased later, by subsequent large-scale impact events.

References:

Nelson et al. (2014) Mapping lunar maria extents and lobate scarps using LROC image products. LPSC Abstract #2861.

Neumann et al. (2015) Lunar impact basins revealed by Gravity Recovery and Interior Laboratory measurements. Science Advances, 1(9), 1-10.

Park et al. (2015) A high resolution spherical harmonic degree 1500 lunar gravity field from the GRAIL mission. AGU Fall Meeting, Abstract #G41B01.

Soderblom et al. (2015) The fractured Moon: Production and saturation of porosity in the lunar highlands from impact cratering, Geophys. Res. Lett., 42, 6939–6944.

Wahl et al. (2020) Crustal porosity of lunar impact basins. Journal of Geophysical Research: Planets, 125, 1-14.

Wieczorek et al. (2013) The crust of the Moon as seen by GRAIL. Science, 339, 671-675.

How to cite: Wahl, D., Wieczorek, M., Wünnemann, K., and Oberst, J.: Porosity Signature of Lunar Impact Basins, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-280, https://doi.org/10.5194/epsc2020-280, 2020.

EPSC2020-305ECP
Tomke Lompa, Kai Wünnemann, Katarina Miljković, Daniel Wahl, and Sebastiano Padovan

Introduction: The lunar surface is characterized by large impact basins that formed in the early history of the Moon. To better understand how basin size and their corresponding gravity signature are related to the impactor size and the thermal history of the Moon, we investigated how properties of impactor and target could affect this formation process. Impactor properties are defined by its size, density, and impactor velocity. The target is characterized by the crustal thickness and the thermal state. Previous studies (e.g.[1],[2],[3],[4]) revealed that thermal conditions in crust and mantle rocks affected the crater formation process, the final basin structure, and the gravity signature. Here, we present numerical simulations for 16 basins located on the lunar farside. The simulations are constrained by gravity data from the Gravity Recovery and Interior Laboratory (GRAIL) mission, and differ in impactor size, crustal thickness and target temperature as a function of depth. We established a relationship between the observed gravity, the geometry of the basin and the pre-impact thermal conditions of the target. These observations lead to a better understanding on how the thermal evolution of the Moon is related to changes in the formation of basins.

Methods: We used the iSALE2D shock physics code to simulate the basin formation process ([5],[6],[7],[8]). We used impactor sizes of 20km to 100km in diameter. The impactor speed was set to 13km/s. The resolution of all models was set to 25 cells per projectile radius (grid cell sizes from 400m to 2000m). The target was composed of a basaltic crust on top of a dunitic mantle. We changed the target properties by using two different crustal thicknesses (40km, 60km) and three temperature profiles which represent the Moon’s thermal states at 4.4, 4.1, and 3.8 billion years ago [9]. Gravity data from GRAIL mission [10] provides detailed information about the deep structure of lunar basins. By combining the most recent GRAIL gravity model GL1500E [11] with LOLA derived topography [12], we obtained a highly resolved Bouguer gravity field.

We assumed constant densities in both crust and mantle to calculate Bouguer gravity anomalies from the basin formation models. This approach can be understood as a simple way to mimic the density distribution after a long-term cooling process in the basin. The shape of the gravity signal preserves information about the position of the crust-mantle boundary and, therefore, assuming constant densities is useful in order to verify the geometry of the model.

Results: For all 54 models we determined the diameter of the largest crustal thickness (DLCT) as a measure of the size of the basins. Figure 1 shows the relationship between impactor sizes and the DLCT for different temperature profiles and initial crustal thicknesses. The results show similar basin diameters for small impactors independently from the thermal profile. For impactor sizes larger than 40km (Fig.1A) or 30km (Fig.1B), the diameters of the hotter profile are significantly higher than those of the cold and intermediate profiles.

We present Hertzsprung and Orientale basin as two examples for gravity modeling. In the formation model of Hertzsprung (Fig. 2A), we used the thermal profile corresponding to an age of 4.1Ga [13]. Figure 2 shows our best-fit model assuming an impactor size of 60km, and a crustal density of 2950kg/m3. The results for Orientale are shown in Figure 2B. In the best-fit model the impactor is 80km in diameter, crustal density is 2750kg/m3 and we assume a cold thermal profile corresponding to an age of 3.8Ga [13]. Gravity profiles from the formation models show that the position of Bouguerminima correlate with the DLCT. The ratio of Bouguerdiameter and DLCT evaluates the fit of the models. A ratio of one indicates that the Bouguerminimum is close to the DLCT.

Conclusion: Our results agree with previous studies [4] in terms of the importance of the thermal state of the target material. Final crater morphologies are sensitive to the thermal conditions in the target material. In our study, temperature effects are visible in simulations with impactors larger than 40km and the change from intermediate to hot thermal states. Isostatic relaxation processes might further modify the initial gravity signature over time. The correlation between basin size and the position of the minimum in observed gravity is a powerful tool to predict impactor diameters without extensive modeling studies.

Acknowledgments: We gratefully acknowledge the developers of iSALE. Topography and gravity data are available at NASA’s PDS. This work is funded by the DFG-grant SFB-TRR 170 (A4). Katarina Miljković research is funded by the Australian Research Council.

References: [1]Ivanov, B., Melosh, H., and Pierazzo, E. (2010) Large Meteorite Impacts and Planet. Evol. IV,29-49. [2]Potter, R. W. K. et al. (2013) JGR:Planets,118,963-979. [3]Zhu, M.-H., Wünnemann, K. and Potter, R. W. K. (2015) JGR:Planets,120,2118-2134. [4]Miljković, K. et al. (2016) JGR:Planets,121,1695-1712. [5]Collins, G. S., Melosh, H. J. and Ivanov, B. A., (2004) Meteoritics & Planet.Sci.,39,217-231. [6]Wünnemann, K., Collins, G. and Melosh, H. (2006) Icarus,180,514- 527. [7]Thompson, S. L. (1990) Sandia National Laboratories Albuquerque, SAND89-2951. [8]Melosh, H. J. (2007) Meteoritics & Planet.Sci.,42,2079–2098. [9]Padovan, S. et al. (2018) EPSC,Abstract#755. [10]Zuber, M. T. et al. (2013) Science,339,668-671. [11]Park, R. S. et al. (2015) AGU Fall Meeting,Abstract#G41-B01. [12]Smith, D. E. et al. (2017) Icarus,283,70-9. [13]Orgel, C. et al. (2017) JGR:Planets,123,1-15.

How to cite: Lompa, T., Wünnemann, K., Miljković, K., Wahl, D., and Padovan, S.: Numerical modeling of farside impact structures on the Moon constrained by gravity data, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-305, https://doi.org/10.5194/epsc2020-305, 2020.

EPSC2020-186ECP
Tiantian Liu, Greg Michael, Wilhelm Zuschneid, Kai Wünnemann, and Jürgen Oberst

1. Introduction

The returned lunar mare soil samples contain highland anorthosites [1]. The source of those non-mare component has been debated since the Apollo epoch ([2]-[9]). In this study, a numerical model was developed to trace the diffusion of non-mare material with the cumulative effect of impact mixing and intermittent filling of mare material.

2. Model

In our previous work [10], a spatially-resolved numerical model was developed to investigate the diffusion of impact melts through impact mixing using the Monte Carlo method. In this study, we extend the capabilities of the model and take into account the volcanic basin infilling forming the lunar mare regions.

2.1 Ejecta distribution

The thickness distribution of proximal ejecta with distance from crater center (r) conforms to a power law, typically: δ(r) = Ar-3 [11]. Since the distal ejecta would fade into the background with continuous impact mixing and the thickness distribution becomes more uniform, we assume that the distal ejecta is also continuously distributed, and the thickness is calculated based on the extrapolated power law. Given the value of the minimum deposit thickness (0.01 m), the ejecta coverage range of each impact event can be calculated.

2.2 Melt volcanic filling

The qualitative eruption features of volcanic history of the Moon were pictured showing a trend of eruption flux from high to low but with different peak times [12]. The thickness distribution of volcanic deposits was systematically studied using crater-geometry techniques pioneered by De Hon (Figure 1b, HDeHon). For each mare region, the areas showing the thickest mare deposit are flooded first. As time goes on, mare material accumulates according to the eruption flux. The surrounding areas of thinner mare deposit are subsequently flooded. After the deposit of the youngest volcanic eruption, the thickness of volcanic deposit corresponds to HDeHon. A time step of 0.01 Ga is used to simulate such continuous eruption. At each step, the mare filling occurred within a certain contour region is taken to be uniformly emplaced (Figure 1a). Volcanic eruptions are assumed to initiate soon after the formation of the respective basin. Due to the great volume and the large flooding areas, only the early eruptions are considered [12]. The rough cessation time of the main filling (tend, Figure 1b) among different mare regions is determined based on the areas of differently-aged mare units.

3. Results and discussions

Given the uncertainty in the flux of mare filling, simulations with the varying setting of mare fillings are run to investigate the major factors affecting the non-mare mixing. Figure 2 shows the results when the mare thickness, cessation time of the major mare filling, and the peak time of eruption are taken to be HDeHon, tend and 3.8 Ga, respectively. By taking an average of the results from multiple simulations where the spatial configuration of generated craters is different, the regional differences are averaged out and the general features of non-mare material concentration are obtained. It shows that throughout the mare regions the non-mare component is present. The older mare surfaces accumulate more non-mare material. If most of the mare regions have filled mare material about 500 m in thickness since the formation of basins, the average and median fractions of non-mare components over all the mare regions are both ~0.1. In the top 5m, the average fraction of the non-mare material within 100 km is both about 0.2; between 100 km and 200 km, the non-mare abundance quickly decreases. In the very top surface (1 m) the non-mare abundance is greater. Apollo 11, 17 and Luna 16 located in the relatively old mare regions have the higher abundance of non-mare components (up to ~0.3). The Apollo 15, 12, and Luna 24 sampling sites located in the relatively young mare regions have a less non-mare component (~0.1).

By comparing the background composition with the geochemical analysis of the lunar mare soil samples, we infer the most plausible geologic processes that have significantly altered the material composition at the sampling sites: for the Apollo 15 and 17 mare soil samples, the large fraction of non-mare material is likely to have resulted from the downslope slumping or lateral transport of the nearby massifs. The Apollo 12 sampling site has a component of Copernicus ejecta. A mixing of both Copernicus ejecta and excavated local underlying material by high-velocity ejecta has altered the composition at the surface. The non-mare material contained in the Apollo 11 and Luna 24 mare soil samples could have been built up gradually by both long-time lateral and vertical mixing. The mare deposit at the Luna 16 landing site is likely to be relatively thin resulting in the abundant local-origin non-mare component.

4. Conclusions

A spatially-resolved numerical model tracing the diffusion of non-mare material was developed to study the source of the non-mare component in the lunar mare soil samples. We find that almost the entirety of the lunar mare regions contains some non-mare components. If the mare thickness is ~500 m, half of the areas possess a fraction ranging from 0.05 to 0.15. By comparing our results with the analysis of the collected soil samples, the most plausible geologic processes that may have altered the non-mare material abundance at the sampling sites are inferred.

References

[1] Wood J.A. (1970), JGR; [2] Rhodes J.M. (1977), Phil Trans Math Phys Eng Sci, 293-301; [3] Simon S.B. et al. (1982), LPSC, 371-388; [4] Huang Y. et al. (2017), JGR-Planets, 1–23; [5] Head J.W. (1982), The moon and the planets, 26(1), 61–88; [6] Budney C.J. and Lucey P.G. (1988), JGR-Planets, 103(E7), 16855–16870; [7] Evan A.J. (2016), GRL, 43(6), 2445–2455. [8] Aridson R. et al. (1975), The Moon, 13(1–3), 67–79; [9] Hörz, F. (1978), LPSC, 3311–3331; [10] Liu T. et al. (2020), Icarus, 113609. [11] Melosh, H.J. (1989), Oxford University. [12] Head J. W. and Wilson L. (1992), Geochim. Cosmochim. Acta, 56(6), 2155–2175.

 

How to cite: Liu, T., Michael, G., Zuschneid, W., Wünnemann, K., and Oberst, J.: Lunar megaregolith mixing by impacts: Evaluation of the non-mare component of mare soils, Europlanet Science Congress 2020, online, 21 Sep–9 Oct 2020, EPSC2020-186, https://doi.org/10.5194/epsc2020-186, 2020.

EPSC2020-592ECP
| MI
Esther Mas-Sanz, Josep M. Trigo-Rodríguez, and Eloy Peña-Asensio

1. Introduction

Lunar meteorites constitute direct, varied samples of the materials forming the lunar surface, allowing to study its bulk chemistry, mineralogy and even geological history. The study of these meteorites not only sheds light on the composition of the Moon but also on its formation and the origin of the Solar System itself. Nowadays we know that these Lunar rocks are delivered to Earth in relative short timescales after being ejected by the impact of an asteroid or a comet from the lunar surface. So far, no lunar meteorite has ever been associated with a crater, not even has been witnessed to fall on Earth [1].

The atmosphere of the Moon is extremely thin, being unable to decelerate particles that are either impacting or escaping from the Moon [2]. It is due to the impact of these high-velocity projectiles that some lunar rocks can be ejected with a velocity above the escape velocity of the Moon (2.38 km/s) [3]. When a projectile of a certain mass impacts the surface of the Moon excavates an impact crater and releases a mass between 3-4 orders of magnitude higher than the projectile mass [3, 4]. Most of the ejected mass falls near the crater and forms continuous ejecta blankets, while the rest follows an impact trajectory far away from the initial impact, while a small part of the ejecta achieves enough velocity to escape the Moon [3, 4]. Our study focuses on the dynamic evolution of these Lunar rocks once they left our satellite, and their ability to reach the Earth as meteorites.

2. Methodology

This study simulates trajectories of lunar ejecta to explore which parameters influence on their fate (impactor/survivor) and their transfer time using the Mercury 6 software [5]. It has been built upon the results and possible time limitations of a previous study [6]. Ejecta trajectories are calculated with a simplified   Solar System model (Earth-Moon-Sun) when time scales are below 1000 years (short transfers). Beyond this value (long transfers, up to 100,000 years), secular perturbations caused by near-by planets are no longer negligible [7] and an 8-planet Solar System model needs to be implemented. Particles found in the vicinity of Earth, at a maximum of 10 Earth's Hill radii, are considered to be in geocentric orbit while those beyond belong to heliocentric orbits. The parameters changed in each simulation were: the local azimuth angle and elevation, launch velocity and the date of the launch. Additionally, the orbital parameters compiled and studied to understand the ejecta evolution and delivery time to our planet.

We have also analyzed a case study corresponding with the impact flare recorded on Sept. 11th, 2013 at 20h07m28s.68 UTC by two telescopes in Seville, Spain (37.34611ºN, 5.98055ºW) provoked by the impact of a meteoroid on the surface of the Moon [8]. The impact was located in the west region of the Mare Nubium, at 17.2ºS 20.5º W (selenographic coordinates). We performed a simulation with 360 particles launched across an arc of 360º from the Mare Nubium impact point, ejected at 45º with respect to the local vertical and with a launch velocity 2.6 km/s. The particle's evolution was then propagated until 104 years to get information about the delivery efficiency, and common timescale.

3. Results

In general we noticed that most ejecta are reaching the Earth in particularly short timescales, between a few days and 10 years. For this fast transfers, predicting if the ejecta will become a future impactor depends only on the launch angle and ejection velocity. Parameters at launch have less and less influence on the fate of the ejecta as time increases. For delivery timescales larger than 1000 years any particle may collide with Earth regardless of its launch angle. The evolution of the main orbital elements for a typical long case is described in Fig. 1. The peak of accumulated collisions shifts to a lower launch velocity at 2.4 km/s.