Magma Oceans in the Inner Solar System

<p>Magma oceans play a fundamental role in the thermal evolution of rocky planets and moons. Most of these objects are thought to experience a magma ocean phase at least at their very early ages. Right after the formation of a planet or moon, its temperatures are expected to be very high mainly due to radiogenic heating from fast-decaying isotopes, and the release of energy due to the gravitational contraction of the object. Therefore, it is very likely that most, if not all, of these bodies, have or have hosted magma oceans. In addition, it is also possible to find, long after their formation, very hot planets or moons that contain large magma oceans in their interior. Perhaps the most famous example is Io, the innermost of the Galilean moons orbiting Jupiter. Due to strong tidal interactions with Jupiter and its moons, Io suffers from extreme tidal heating, strong enough to melt large regions of its interior.</p> <p>On the other hand, mantle convection simulations are crucial to understanding the long-term thermal evolution of rocky planets and moons. These simulations are tailored to solid convection in the mantle, thus when considerable fractions of melt are present or are generated in the interior of the planet, the contrasting properties of the solidus and liquidus phases become too large for the simulations to handle. Properties such as the viscosity and the thermal conductivity are several orders of magnitude different between solidus and liquidus. As consequence, the time steps of the simulation become too short, and the simulation ends up 'frozen' in time. Besides, these large differences pose numerical restrictions that make the codes unstable and prone to crash. Due to these limitations, simulating the interior thermal evolution of hot rocky planets and moons with mantle convection codes turns out to be extremely challenging.</p> <p>In this work, we apply our magma oceans modeling to Io, since it is an appropriate test case in the Solar System with evidence of hosting large amounts of magma in its interior (e.g., Khurana et al. 2011). We use the CHIC convective code (Noack et al. 2015) to model Io's interior thermal evolution. We then apply and expand an approach, originally developed by Golabek et al. 2011, in which the effective local properties are altered wherever melt is present or produced. Our modeling can be applied to any rocky planet with hot interiors, as could be the case of the innermost Trappist-1 planets, or many of the rocky exoplanets known.</p>

<sec><p indent="0mm">The Earth is distinguished from other planets by its unique plate tectonics and the presence of an evolved felsic continental crust. The growth of the continental crust is not uniform, with most of it forming within the first 1.5 billion years of Earth’s history (i.e., before ~3.0 Ga). The preserved Archean continental crust is primarily composed of greenstones and felsic rocks such as tonalite-trondhjemite-granodiorite (TTG), many of which were formed between 3.0 and 4.0 Ga or even earlier. Global plate tectonics likely began in the mid-Archean, raising the question of whether the earliest felsic rocks were formed in localized plate tectonic settings before the initiation of global plate tectonics or through alternative, non-plate tectonic mechanisms. Previous models suggest the earliest TTG formation through partial melting at the base of oceanic plateaus under a stagnant lid regime, but these fail to explain the source of water required for the melting of basaltic rocks. </sec><sec> The early Earth experienced extensive melting, forming a global magma ocean. Once Earth cooled sufficiently for heat conduction to no longer sustain surface melt temperatures, the silicate melt inevitably solidified, forming a crust. Even with a thin solid crust forming on the surface of the magma ocean, heat could still be effectively dissipated through conduction and radiation. The cooling rate of the magma ocean, though rapid on a geological timescale, was quite slow compared to the crystallization of silicate minerals in basaltic melts. Consequently, fractional crystallization of olivine in the upper magma ocean could have led to melt differentiation from ultramafic to mafic, forming a mafic proto-crust. </sec><sec> The formation and thickness of the mafic proto-crust significantly altered the way and efficiency of volatile components, such as water, migrating during the degassing process of the magma ocean. As the proto-crust thickened, its barrier effect caused the degassing efficiency of the magma ocean to plummet, reducing the rate of magma degassing and ultimately approaching zero. Due to the extremely high incompatibility of water and other volatile components within solid substances, once a solid crust forms on the surface of a magma ocean, these components will be expelled from the solid material and enter the residual magma ocean during the solidification process. Therefore, before the magma ocean completely solidifies, most of the H₂O will remain in the residual melt. </sec><sec> This outward-to-inward cooling and solidification progressively thickened the crust from the top down. As the proto-crust thickened, the solid-melt interface migrated downward, driving water and other volatiles toward the interior. This process, combined with the solid mantle accretion from deep to shallow regions, would have concentrated water in the residual melt, ultimately forming a water-rich layer at the base of the proto-crust. Therefore, during the cooling of the magma ocean, the “reverse migration” of water due to the thickening of the proto-crust may have sequestered a large amount of water in the solid Earth, creating conditions for the hydrous melting of the proto-crust to form the earliest TTG. </sec><sec> This mechanism not only explains the high water content required for basaltic proto-crust melting to form felsic magmas before the onset of plate tectonics but also accounts for the observed transition in the formation pressure of Hadean and Eoarchean sodic granitoids from low to high over time. Previous researchers found that a significant change in Archean sodic granitoid composition occurs in the 3.9–3.75 Ga period, corresponding to an increase in formation depth from <sc>~3 km</sc> to <sc>25-50 km.</sc> We hypothesize that the hydrous melting triggered by top-down water enrichment caused by the proto-crust thickening may have mainly occurred during the Hadean and the Eoarchean. </sec>

The lunar anorthosite highlands, composed of 90 wt% anorthite (Ca-rich plagioclase), represent the Moon&amp;#8217;s primary crust [1], which formed during the solidification of a magma ocean following the Moon-forming giant impact between the proto-Earth and Theia. The composition of this anorthite-enriched crust is unique among the terrestrial bodies within our Solar System and can be easily reproduced by crustal formation due to the flotation of buoyant anorthite crystals during magma ocean solidification [2] (Fig. 1a). However, this canonical model often struggles to reproduce the long, &gt; 200 Myr solidification timescale required by the ages of the anorthosite suite [2], as well as the diverse magmatic histories of the lunar anorthosites [3].If, instead, the anorthite crystals remain entrained in magma ocean due to vigorous convection, the magma ocean behaves as a liquid-crystal slurry with an immobile stagnant lid at the surface (Fig. 1b), which reduces the cooling rate of the magma ocean and results in long solidification durations [3]. In this regime, crustal formation occurs via extraction of melt from the magma ocean through the stagnant lid. In this study, we investigate whether sufficient magmatic differentiation can occur within the stagnant lid to produce an anorthite-rich crust. Our model consists of a simplified two-component anorthite-olivine melt system, which becomes progressively enriched in anorthite as the melt solidifies. We calculate the thermochemical evolution of this melt as it rises through a growing stagnant lid, cools, and solidifies for a range of Peclet numbers (the ratio of the rate of advection to diffusion), and lid thicknesses.We find that for low Peclet numbers (30), thick, melt- and anorthite-rich (&gt;30 vol%) sills can form within the lid and migrate slowly upwards. However, the composition of these sills is still more olivine-rich than that observed in the lunar highlands. Therefore, further differentiation is likely required within the sills themselves. This could occur by the flotation of anorthite crystals as this is a process that is more likely to occur in these emplaced magmatic bodies than at the scale of the magma ocean. Equally none of these sills are predicted to reach the cold, solid surface of the lid. However, we find that the overpressure within the sills is sufficient to drive the eruption of this melt to the surface, leading to a crustal formation by a combination of extrusive and intrusive magmatism.[1] Wood et al., 1970, in Proceedings of the Apollo 11 Lunar Science Conference (pp. 1965-1988)[2] Elkins-Tanton, L.T., Burgess, S. and Yin, Q.Z., 2011. The lunar magma ocean: Reconciling the solidification process with lunar petrology and geochronology. Earth and Planetary Science Letters, 304(3-4), pp.326-336.[3] Ashwal, L.D. and Bybee, G.M., 2017. Crustal evolution and the temporality of anorthosites. Earth-Science Reviews, 173, pp.307-330.[4] Michaut, C. and Neufeld, J.A., 2022. Formation of the lunar primary crust from a long&amp;#8208;lived slushy magma ocean.&amp;#160;Geophysical Research Letters,&amp;#160;49(2), p.e2021GL095408.Fig. 1. Sketches of a) the classical anorthite flotation crust model and b) the slushy magma ocean model where the crystal fraction remains entrained by convection and a crust forms by melt extraction in a stagnant lid.

Convective outgassing efficiency in planetary magma oceans: Insights from computational fluid dynamics

&amp;lt;p&amp;gt;Terrestrial planets evolve through various stages of large-scale melting, or magma oceans, due to the energy release during accretion and differentiation. Any magma ocean is thought to become progressively enriched in FeO and incompatible elements upon freezing due to fractional crystallization. The resulting upwards enrichment of the related cumulate (=crystal) packages drives gravitational overturn(s) of the incipient mantle, and ultimately stabilizes a FeO-enriched molten layer at the core-mantle-boundary (CMB)&amp;lt;sup&amp;gt;1&amp;lt;/sup&amp;gt;. Such a molten layer, previously termed basal magma ocean (BMO)&amp;lt;sup&amp;gt;2&amp;lt;/sup&amp;gt;, is thought to also fractionally crystallize, but downwards instead of upwards, and over much longer timescales than the surficial magma ocean. This BMO fractional crystallization due to slow planetary cooling analogously implies the stabilization of a thick FeO-enriched layer at the CMB. Such a layer would essentially remain stable forever, as being too dense to be entrained by convection of the overlying mantle. However, at least for Earth, geophysical observations rule out the preservation of such a deep dense global layer. Here, we investigate the consequences of an alternative mechanism for BMO freezing, reactive crystallization, on the initial condition of solid-state mantle convection and long-term planetary evolution.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;Based on scaling relationships, we show that any cumulates, which crystallize from the BMO (e.g., due to initial cooling or reaction) are readily entrained by mantle convection. Once the BMO-mantle boundary is exposed, the BMO reacts with the mantle to form reactive cumulates. Reaction is driven by disequilibrium between mantle rocks and the BMO, a situation that is inevitable independent of BMO initial composition. As reactive cumulates are continuously entrained by mantle convection, the BMO continues to freeze by reactive crystallization. Based on lower-mantle mineral-melt phase equilibria&amp;lt;sup&amp;gt;3&amp;lt;/sup&amp;gt;, we calculate the compositional evolution of the BMO, and the chemistry of the BMO cumulate package. We demonstrate that for a wide range of BMO initial compositions, the cumulate package consists of two discrete layers: the first is pure bridgmanite close to the MgSiO&amp;lt;sub&amp;gt;3&amp;lt;/sub&amp;gt; end-member; the second is mostly bridgmanite+ferropericlase that is moderately enriched in FeO and incompatibles, i.e. similar in composition to FeO-enriched pyrolite. The mass or thickness of the cumulate package depends on reaction kinetics, but is significantly larger than that of the BMO. The bridgmanitic layer is expected to be entrained by mantle convection due to its intrinsic buoyancy, but resist efficient mixing due to its intrinsic strength, thereby potentially providing an explanation for seismic scatterers/reflectors and ancient geochemical reservoirs&amp;lt;sup&amp;gt;4&amp;lt;/sup&amp;gt;. The moderately FeO-enriched layer is expected to stabilize thermochemical piles, providing a candidate origin for the seismically-observed large low shear velocity provinces (LLSVPs)&amp;lt;sup&amp;gt;5&amp;lt;/sup&amp;gt;.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;These results have implications for the long-term (thermal) evolution of planets in general. Earth-sized terrestrial (exo-)planets and super-Earths should also initially host a MgSiO&amp;lt;sub&amp;gt;3&amp;lt;/sub&amp;gt;-rich layer as well as a moderately FeO-enriched layer. In contrast, small terrestrial planets such as Mars may host a more strongly Fe-rich deep dense global layer as long as no BMO is stabilized in their histories.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;[1] Ballmer+, G-cubed 2017; [2] Labrosse+, Nature 2007; [3] Boukar&amp;amp;#233;+, JGR Solid-Earth 2015; [4] Ballmer+, Nat.Geosci. 2017; [5] Ballmer+, G-cubed 2016.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;&amp;amp;#160;&amp;lt;/p&amp;gt;

Context. A magma ocean (MO) is thought to be a ubiquitous stage in the early evolution of rocky planets and exoplanets. During the lifetime of the MO, exchanges between the interior and exterior envelopes of the planet are very efficient. In particular, volatile elements that initially are contained in the solid part of the planet can be released and form a secondary outgassed atmosphere. Aims. We determine trends in the H–C–N–O–S composition and thickness of these secondary atmospheres for varying planetary sizes and MO extents, and the oxygen fugacity of MOs, which provides the main control for the atmospheric chemistry. Methods. We used a model with coupled chemical gas-gas and silicate melt-gas equilibria and mass conservation to predict the composition of an atmosphere at equilibrium with the MO depending on the planet size and the extent and redox state of the MO. We used a self-consistent mass–radius model for the rocky core to inform the structure of the planet, which we combined with an atmosphere model to predict the transit radius of lava worlds. Results. The resulting MOs have potential temperatures ranging from 1415 to 4229 K, and their outgassed atmospheres have total pressures from 3.3 to 768 bar. We find that MOs (especially the shallow ones) on small planets are generally more reduced, and are thus dominated by H2-rich atmospheres (whose outgassing is strengthened at low planetary mass), while larger planets and deeper MOs vary from CO to CO2–N2–SO2 atmospheres, with increasing $\[f_{\mathrm{O}_2}\]$. In the former case, the low molecular mass of the atmosphere combined with the low gravity of the planets yields a large vertical extension of the atmosphere, while in the latter cases, secondary outgassed atmospheres on super-Earths are likely significantly shrunk. Both N and C are largely outgassed regardless of the conditions, while the S and H outgassing is strongly dependent on the $\[f_{\mathrm{O}_2}\]$, as well as on the planetary mass and MO extent for the latter. We further use these results to assess how much a secondary outgassed atmosphere may alter the mass–radius relations of rocky exoplanets.

&amp;lt;p&amp;gt;Rocky planets go through at least one and likely multiple magma ocean stages, produced by the giant impacts of accretion. Planetary data and models show that giant impacts do not dehydrate either the mantle or the atmosphere of their target planets. The magma ocean liquid consists of melted target material and melted impactor, and so will be dominated by silicate melt, and also contain dissolved volatiles including water, carbon, and sulfur compounds.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;As the magma ocean cools and solidifies, water and other volatiles will be incorporated into the nominally anhydrous mantle phases up to their saturation limits, and will otherwise be enriched in the remaining, evolving magma ocean liquids. The water content of the resulting cumulate mantle is therefore the sum of the traces in the mineral grains, and any water in trapped interstitial liquids. That trapped liquid fraction may in fact be by far the largest contributor to the cumulate water budget.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;The water and other dissolved volatiles in the evolving liquids may quickly reach the saturation limit of magmas near the surface, where pressure is low, but degassing the magma ocean is likely more difficult than has been assumed in some of our models. To degas into the atmosphere, the gases must exsolve from the liquid and form bubbles, and those bubbles must be able to rise quickly enough to avoid being dragged down by convection and re-dissolved at higher pressures. If bubbles are buoyant enough (that is, large enough) to decouple from flow and rise, then they are also dynamically unstable and liable to be torn into smaller bubbles and re-entrained. This conundrum led to the hypothesis that volatiles do not significantly degas until a high level of supersaturation is reached, and the bubbles form a buoyant layer and rise in diapirs in a continuum dynamics sense. This late degassing would have the twin effects of increasing the water content of the cumulates, and of speeding up cooling and solidification of the planet.&amp;lt;/p&amp;gt;&amp;lt;p&amp;gt;Once the mantle is solidified, the timeclock until the start of plate tectonics begins. Modern plate tectonics is thought to rely on water to lower the viscosity of the asthenosphere, but plate tectonics is also thought to be the process by which water is brought into the mantle. Magma ocean solidification, however, offers two relevant processes. First, following solidification the cumulate mantle is gravitationally unstable and overturns to stability, carrying water-bearing minerals from the upper mantle through the transition zone and into the lower mantle. Upon converting to lower-mantle phases, these minerals will release their excess water, since lower mantle phases have lower saturation limits, thus fluxing the upper mantle with water. Second, the mantle will be near its solidus temperature still, and thus its viscosity will be naturally low. When fluxed with excess water, the upper mantle would be expected to form a low degree melt, which if voluminous enough with rise to help form the earliest crust, and if of very low degree, will further reduce the viscosity of the asthenosphere.&amp;lt;/p&amp;gt;

Mechanisms of metal–silicate equilibration in the terrestrial magma ocean

Resolving early silicate differentiation timescales is crucial for understanding the chemical evolution and thermal histories of terrestrial planets. Planetary-scale magma oceans are thought to have formed during early stages of differentiation, but the longevity of such magma oceans is poorly constrained. In Mars, the absence of vigorous convection and plate tectonics has limited the scale of compositional mixing within its interior, thus preserving the early stages of planetary differentiation. The SNC (Shergotty-Nakhla-Chassigny) meteorites from Mars retain 'memory' of these events. Here we apply the short-lived 146Sm-142Nd and the long-lived 147Sm-143Nd chronometers to a suite of shergottites to unravel the history of early silicate differentiation in Mars. Our data are best explained by progressive crystallization of a magma ocean with a duration of approximately 100 million years after core formation. This prolonged solidification requires the existence of a primitive thick atmosphere on Mars that reduces the cooling rate of the interior.

Similar to other terrestrial planets like Moon and Mars, Earth experienced a magma ocean period about 4.5 billion years ago. On Earth differentiation processes in the magma ocean set the initial conditions for core formation and mantle evolution. During the magma ocean period Earth was rotating significantly faster than today. Further, the viscosity of the magma was low, thus that planetary rotation potentially played an important role for differentiation. However, nearly all previous studies neglect rotational effects. All in all, our results suggest that planetary rotation plays an important role for magma ocean crystallization. We employ a 3‐D numerical model to study crystal settling in a rotating and vigorously convecting early magma ocean. We show that crystal settling in a terrestrial magma ocean is crucially affected by latitude as well as by rotational strength and crystal density. Due to rotation an inhomogeneous accumulation of crystals during magma ocean solidification with a distinct crystal settling between pole and equator could occur. One could speculate that this may have potentially strong effects on the magma ocean solidification time and the early mantle composition. It could support the development of a basal magma ocean and the formation of anomalies at the core‐mantle boundary in the equatorial region, reaching back to the time of magma ocean solidification.

Settling of metal droplets in a terrestrial magma ocean: On the correction of the Stokes velocity

Onset of magma ocean solidification on Mars inferred from Mn-Cr chronometry

Lava worlds: From early earth to exoplanets

The split fate of the early Earth, Mars, Venus, and Moon

Introduction: The Ambiguity of the Term “Magma Ocean” Applied to Asteroids The suggestion that Earth’s Moon (e.g. Warren, 1985; Taylor and Norman, 1991) and Earth itself (e.g. Agee and Longhi, 1992) had magma oceans in their very early histories, coupled with the realization that 26 Al present in early-formed asteroids provided a heat source for partial and perhaps total melting (e.g. Hevey and Sanders, 2006; McCoy et al., 2006a; Wadhwa et al., 2006; Schiller et al., 2010), prompted study of the magma ocean concept on asteroids. However, this use of the term magma ocean in relation to melting in asteroid mantles is very different from its original use to describe the molten outer part of the Moon immediately after its formation. The Moon is commonly inferred to have experienced melting of its outer several hundred kilometers as a result of its very rapid accretion from material left in Earth’s orbit after the impact of a Mars-sized body on the proto-Earth (Canup, 2004), though many geochemical and isotopic problems with this scenario have emerged (Asphaug, 2014). Other terrestrial planets may also have had near-surface magma oceans, produced by less frequent but higher-energy impacts of planetesimals that were the last contributors to the growth of these bodies (e.g. Rubie et al., 2015). In all these cases, the magma ocean would have formed all of the outer part of the body and covered a much cooler interior. In contrast, asteroid magma oceans, of the kind currently generally implied in the literature, would have been produced in the interior of bodies by the accumulation of heat from the decay of, predominantly, the short-lived but very energetically radioactive isotope 26 Al (e.g. McCoy et al., 2006a), present at the time of solar system formation. If an asteroid formed sufficiently early, its interior could have reached a temperature well above the solidus of all common minerals (Hevey and Sanders, 2006; McCoy et al., 2006a; Wadhwa et al., 2006; Schiller et al., 2010). If melts did not migrate after formation, a very large fraction of partial melting - up to complete silicate melting - below an outer, cool crustal layer is theoretically possible. Such a body of liquid silicate would clearly qualify as a subcrustal magma ocean.