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Models of hydrothermal activity on icy satellites
One question in evaluating the potential habitability of other worlds is the length of time that they can support habitable environments, such as hydrothermal vents, given the constraints detailed above. The emergence of life is sometimes framed as a probability based on a number of favorable conditions (Forgan, 2009), many of which are inherently undeterminable as long as the true origins of life are still under debate. However, as specific environmental conditions on geologically active worlds are by definition transient, the temporal factor at least can be considered (Ćirković, 2004). If favorable conditions for the emergence of life are only available for a vanishingly short time, then the likelihood of life emerging there is similarly vanishingly small. Defining the timescales of these environmental conditions on icy satellites of course presents its own challenges. Even the existence of hydrothermal vents in Enceladus is indirectly inferred from few, though exciting, data. Evaluating the geologic histories of these satellites, including the presence or duration of liquid oceans or hydrothermal activity, has then long relied on modelling constrained by remote observations.
In the Saturnian system, orbital and geophysical models have progressed rapidly in the past 20 years with the observational data provided by the Cassini mission. In particular, new geophysical constraints on the densities and differentiation within each moon have spurred an active debate on the ages and orbital histories of Saturn’s moons and rings. The geophysical evidence for a liquid ocean on Enceladus that completely separates its icy crust from a rocky inner core has superseded previous assertions that the moon is too small to sustain a tidally heated liquid ocean. This, in turn, inspired new models of tidal heating (Choblet et al., 2017; Liao et al., 2020), which again must be reconciled with the evolution of the other moons. Thermal evolution models of Saturn’s innermost moons show that tidal heating could sustain a liquid ocean on Enceladus for up to 1 Gyr, supposing a 4.5 Gyr orbital history (Neveu & Rhoden, 2019; Fig. 1.3). Other lines of reasoning have proposed radically different histories that attain the same current state of Enceladus, such as the analysis of Ćuk et al. (2016) that proposed a recent (< 100 Mya) disruption and re-accretion of Saturn’s inner moons in order to explain their current orbital configurations. In short, while observations from the Cassini mission have provided a new wealth of geophysical constraints, these constraints have not yet led to a consensus on the ages, orbital histories, or structural evolutions of Saturn’s moons.
Fig. 1.3: Thermal evolution of Enceladus (modified from Neveu and Rhoden, 2019). The legend indicates color for temperatures -273 to 327 ⁰C (0-600 K), with a demarcation at the liquidus (total melting point, at 0 ⁰C) of the NH3-H2O system. Y axis depths range from core (0) to surface. Yellow arrow at “Ocean” indicates the duration of Enceladus’ history where the temperature inside the moon is > 0 ⁰C and a liquid water ocean differentiates from the rocky core.
The recent and current scientific discussions exploring the history and current state of water-rock interaction within Enceladus is similarly active. Based on the H2:H2O ratio in Enceladus’ plumes, Waite et al. (2017) calculated the H2 release rate from Enceladus as 109 moles yr-1, the major proportion of which is interpreted to be attributable to inorganic mineral alteration. Based on the total amount of H2 producible from a rock core the size of Enceladus’, water-rock alteration could sustain this release rate for hundreds of millions of years. This can be compared to other models of water-rock interaction in Enceladus’ core. The 1D reaction models of (Malamud & Prialnik, 2013, 2016) evaluate heat generation by water transport through the core, including a fixed alteration rate for primary minerals. As they do not assume a mechanism for the re-introduction of water into the porous core, their results show that water forced out of the core by compaction pressures quickly freezes, forming an ice crust in contact with the rock core before the primary mineralogy can be completely altered to secondary hydrous minerals. However, their results also show that even during this relatively brief geologic period (~20 Myr), the majority of primary minerals within the core are converted to hydrous secondary phases. Similarly, fracture-based models of primary mineral alteration by water do not yield results that extend the alteration timespan (Vance & Melwani Daswani, 2020).
Fig. 1.4: a. conceptual model the interior of Enceladus and b. an associated model of water flow and water-rock interaction regimes at Enceladus’ shallow vs. deep core, as presented by Glein et al. (2018). The water-rock interaction regimes describe a differentiated core where carbonate geochemistry dominates the shallow core and serpentinized rock in a reduced environment is the major component of the deep core.
The existence and relative abundances of salts, N2, NH3, CO2, CH4, H2, and silica nanoparticles associated with the water vapor emitted from Enceladus have additionally been used to constrain various properties of Enceladus’ liquid ocean and possible water-rock reactions therein. The CO2:CH4 ratios of the plume have been calculated to indicate largely reducing redox conditions (Glein et al., 2008). The presence of NH3 and the calculated carbonate stability indicate an alkaline pH (Glein et al., 2015; Hsu et al., 2015) that is constrained to pH 8.5-10.5 (for a fluid at 25 ⁰C), based on experimental and modelled silica nanoparticle formation and stability (Hsu et al., 2015; Sekine et al., 2015). Recent models of seafloor hydrothermal vents in Enceladus explore more localized geochemical scenarios, postulating a general divide between shallow, carbonate-driven seafloor water-rock interactions and deep-core alteration dominated by serpentinization reactions (Glein et al., 2018; Glein & Waite, 2020; Ray et al., 2020). Based on equilibrium models of the geochemical co-evolution of a rocky core having an idealized chondritic mineralogy and circulating water, Zolotov (2007) proposed a range of geochemical compositions for a liquid reservoir as might exist in Enceladus. The range proposed by Zolotov’s model encompasses the range of ocean compositions subsequently calculated from the salts found in Enceladus’ emissions (Waite et al., 2009, 2017), although some constituents of the model (namely magnesium and sulfur species) have not been identified in the plumes or E-ring and as such cannot be evaluated by comparison. All these models, consistent with the geochemical data collected by Cassini, indicate a history of hydrothermal water-rock interaction in Enceladus that has resulted in an ocean approaching equilibrium with a chondritic rocky core.
Experimental insights into olivine alteration on icy satellites
In any discussion of serpentinization-driven hydrothermal activity, the mechanism of serpentinization itself is a critical focus of inquiry. While the mechanisms between elementary reactions and overall reactions (usage as defined by Lasaga, 1984) are intimately interrelated, experimental studies by design often focus on fundamental processes by isolating reactions to better understand natural processes. For the purposes of this discussion, studies of olivine alteration in aqueous fluids can be loosely grouped by their focus on either olivine dissolution, or serpentinization. Dissolution studies as discussed here are generally structured to attain a steady-state, stoichiometric olivine dissolution rate. Most studies also provide information on the kinetic factors that influence measured dissolution rates. The serpentinization studies considered here, in a general sense, focus on the hydration of olivine or ultramafic rocks, and in particular the rates and mechanisms of the entire conversion process from primary to secondary phases.
Olivine dissolution
San Carlos olivine (Mg1.83Fe0.17SiO2, which is close in composition to the pure Mg-end-member mineral forsterite) provides a convenient model mineral for evaluating alteration processes that may occur throughout the solar system, as magnesium-rich olivine is a common constituent of chondrites and meteorites. San Carlos olivine dissolution can be described by the following reaction: Mg1.82Fe0.17SiO4 + 4 H+ → 1.82 Mg2+ + 0.17 Fe2+ + SiO2 + 2 H2O Equation 1.1
The study of forsterite-water interactions benefits from decades of experimental research that report the rates and mechanisms of olivine dissolution (as well as overall serpentinization reactions, discussed below). Olivine dissolution as in Eq. 1.1 can be thought of as an elementary reaction of complex serpentinization reactions (although in reality, the process of dissolution itself can be framed as a complex reaction where breaking and rearranging each of the Mg-O and SiO-Mg bonds are elementary reaction steps. See for example Rimstidt et al. (2012) for a review).
Empirical studies of San Carlos olivine show that the strongest determining factors of the dissolution rate are temperature and pH. At low temperatures and pressures relevant to Enceladus (0-150 ⁰C, up to 400 bar), the dissolution rate of olivine is one of the fastest of all silicate minerals, in accord with the Goldich mineral stability series formulated in the 1930s (Goldich, 1938). Some studies have shown that olivine dissolution is as much, or even more, sensitive to changes in pH as compared to temperature in the 0-150 ⁰C range (Hänchen et al., 2006). Like other Mg-silicates (Oelkers, 2001), the solubility of olivine is generally modelled as decreasing with pH increase, even in very high pH fluids (pH 11-14). Based on aggregated empirical findings, the general mechanism of olivine alteration is interpreted to change from acid to alkaline conditions (e.g. Crundwell, 2014; Palandri & Kharaka, 2004; Pokrovsky & Schott, 2000a, 2000b). This generally assumes two parallel reactions: one at silica-rich surface sites that is active in acidic conditions and another at hydrated Mg surface sites that dominates in alkaline conditions (Fig. 1.5a). This conceptual model for forsterite dissolution is expressed as follows (as from Pokrovsky & Schott, 2000b): R = kSi {>Si2O-H+] + kMg {>MgOH2+} Equation 1.2
In the above equation, R indicates the overall dissolution rate in mol cm-2 s-1, ki indicates a rate constant for Si and Mg and {>j} designates the concentration of surface species j in mol m-2. The output of this empirically derived model describes a generally weaker dependence of dissolution rate on pH at alkaline vs. acid conditions. This is borne out by several recent meta-analyses of olivine dissolution rate studies (Brantley, 2003; Oelkers et al., 2018; Rimstidt et al., 2012).
Fig. 1.5 Forsterite dissolution mechanisms and rates. a. Forsterite dissolution mechanism in acid vs. alkaline solutions, adapted from Pokrovsky and Schott (2000b). b. Aggregated dissolution rate equations of olivine dissolution over pH, as presented by Oelkers et al. (2018). Lines plot the rate equation given by each of the noted publications (Crundwell, 2014; Palandri & Kharaka, 2004; Pokrovsky & Schott, 2000b; Rimstidt et al., 2012; Wogelius & Walther, 1991) and the grey shaded area indicates the approximate limits of the experimental data used in generating the plotted rate equations.
Olivine dissolution studies are notably better represented in the low- to neutral-pH range, and as pH increases the variance between calculated rates also increases (grey shaded area in Fig. 1.5b). Throughout the entire pH range, olivine dissolution often starts out non-stoichiometric before converging to a stoichiometric steady-state dissolution rate: at acid-to-neutral pH solutions, the initial behavior is Mg > Si, with stoichiometric steady-state dissolution occurring after several hours; in alkaline pH solutions, initially Si > Mg, with stoichiometric dissolution being achieved only after days or weeks (Pokrovsky & Schott, 2000b). Rates calculated from data obtained during the initial period do not therefore reflect the stoichiometric steady-state dissolution rate. The inclusion of rates calculated during experiments at non-stoichiometric dissolution may in part be the reason for the increased variance between rate data in alkaline conditions. In addition, the ambiguity of whether rates reported by all studies represent stoichiometric dissolution is far from the only potential difficulty in comparing rate data between different experimental data sets. Select studies of olivine alteration have shown the formation of an amorphous silica layer on the olivine surface that may slow olivine dissolution (Béarat et al., 2006; Sissmann et al., 2013), and that may even start forming within a matter of days (Daval et al., 2011). These same types of surface altered layers (SAL) have been found in olivine reacted in flow-through experiments, implying that surface interfacial processes control the formation of SALs even when they are not predicted to form by bulk fluid chemistry (Hellmann et al., 2012). This calls into question the “steady-state” assumption of “stoichiometric steady-state” dissolution, as it implies that this apparent steady-state dissolution rate may actually decrease over time. To date, detailed studies at the nm-scale of fluid-solid interfaces have been restricted to olivine reacted in acid solutions; at basic pH conditions, there is a distinct lack of data on surface alteration reactions.
Serpentinization
Serpentinization is a process closely related to olivine dissolution that was originally formulated to describe the exothermic conversion of ultramafic primary minerals (namely, olivine and pyroxene) to serpentine minerals (chrysotile, lizardite, and/or antigorite), magnetite, and H2. The serpentinization of olivine specifically can be described by the following reactions:
2 Mg2SiO4 (forsterite) + 3 H2O → Mg3Si2O5(OH)4 (serpentine) + Mg(OH)2 (brucite)
3 Fe2SiO4 (fayalite) + 2 H2O → 2 Fe3O4 (magnetite) + 3 SiO2 + 2 H2
As with olivine dissolution, rates of serpentinization are similarly sensitive to temperature and pH. An abundance of studies constrain serpentinization at 200-350 ⁰C, temperatures relevant to mid ocean ridge hydrothermal systems (Elderfield & Schultz, 1996). Generally, the rate of serpentinization increases with temperature, with maximum serpentinization rates occurring over a range of temperatures (250-310 ⁰C) (Lamadrid et al., 2020; Malvoisin et al., 2012; Wegner & Ernst, 1983). The rate of conversion from primary to secondary minerals should not be considered to be a simple process (Seyfried et al., 2007), an assertion that is perhaps evident from the fact that “serpentinization” can start with a single mineral, or a complex ultramafic mineral assemblage. In addition, fluid chemistry properties such as pH, ionic strength, and redox potential often dynamically change during experiments. These complexities, of course, can then affect the rate of formation and stability of any secondary phases (Janecky & Seyfried, 1986). For example, (McCollom et al., 2020) recently showed that the type and abundance of secondary phases is sensitive to both the initial mineral assemblage (olivine vs. olivine-orthopyroxene) and to pH, while (Tutolo et al., 2018) found that the silicification of secondary brucite into tertiary serpentine can act as a rate-limiting step in serpentinization at 150 ⁰C. Given these complexities, it is not surprising that reported rates of serpentinization vary by up to 5 orders of magnitude (Andreani et al., 2013; Lafay et al., 2012; Malvoisin et al., 2012; Ogasawara et al., 2013; Wegner & Ernst, 1983). All of these aspects of serpentinization rates and mechanisms have been recently reviewed by McCollom et al. (2016).
The heat from exothermic serpentinization reactions and the H2 generated by Fe(II) oxidation to Fe(III) in magnetite (Eq. 1.4) are of particular interest to habitability constraints, as they may provide thermal and chemical energy where solar energy is nonexistent (Truche et al., 2020). While there remains no consensus on the lower temperature limits of serpentinization and little data to constrain rates of serpentinization at low temperatures, a number of recent studies have led to progress in this field (Table 1). Field studies have identified serpentine minerals and H2 that are thought to have been produced at in-situ temperatures as low as 30 ⁰C (H. M. Miller et al., 2016; Okland et al., 2012). Magnetite formation (Eq. 1.4) has been identified during low-temperature alteration experiments (H. M. Miller et al., 2017), but even where its formation is inhibited, iron oxidation can occur due to Fe(III) incorporation into serpentine, talc, and Fe(III)(oxy)-hydroxides (Mayhew et al., 2013; Neubeck et al., 2014; Seyfried et al., 2007). Low-T alteration experiments have been shown to produce H2 from mafic mixed-mineral assemblages as low as 50 ⁰C (Neubeck et al., 2016), and monomineralic low-temperature studies of San Carlos olivine alteration have shown H2 production at a range of temperatures as low as 25 ⁰C in pure water (Okland et al., 2014) and 50 ⁰C in carbonated water (Neubeck et al., 2016). Motivated by identifying alteration products on Martian surfaces, experiments focusing on secondary products have pushed the temperature limits even further. Olivine and basalt dissolution in highly acidic, unfrozen brines has been shown to continue to -19 ⁰C (Hausrath & Brantley, 2010), and sulfates were identified after olivine alteration in partially frozen sulfuric acid solutions at -40 and -60 ⁰C (Niles et al., 2017).
NH3 effect on olivine alteration
The evolution of hydrothermal vents on ocean worlds may still be quite distinct from terrestrial environments. Icy satellites beyond the frost line in the solar system in some cases never experience high temperatures, and so can preserve volatile species such as NH3 and methanol that are miscible in water (Kargel, 1992; Waite et al., 2009). These volatile species may affect the evolution of hydrothermal vents by affecting the dissolution of primary minerals, or by affecting the serpentinization processes by stabilizing or otherwise controlling secondary phases. Considering the potential effect of NH3 on olivine dissolution, a review of rate limiting factors on olivine dissolution suggest that changes to water activity, density, or nitrogen content are not known to significantly affect olivine dissolution rates (Rimstidt et al., 2012). However, NH4+ conspicuously resembles H2O in several respects (molar masses, bond angles, interatomic distances, partial molar volumes, and hydrogen bond strengths) and may therefore stabilize dissolved species owing to its tetrahedral shape and its ability to form hydrogen bonds (Brugé et al., 1999; Perrin & Gipe, 1987), or via redox processes (Chivers & Elder, 2013). The effects of these parameters on the rate constants of olivine dissolution, by potentially interacting with reactant species at the mineral surface or in the fluid, have not explicitly been explored. Some studies have already shown that NH3 can stabilize minerals that are also secondary products of olivine alteration: industrial processes have proposed NH3 as a process step in managing grain size distributions in silica nanoparticle production (Lazaro et al., 2013; Raza et al., 2018; Wang et al., 2011) and, in conjunction with carbonate species, in stabilizing metastable magnesium carbonate minerals such as nesquehonite and roguinite (Zhu et al., 2017). This in turn suggests the possibility that NH3 may affect primary mineral dissolution rates by controlling the production and stability of secondary minerals formed at the primary mineral surface, inhibiting primary mineral interaction with the bulk water phase.
Synthesis and alteration of organics in exotic hydrothermal environments
Abiotic organic synthesis and concentration is a key point in discussions of the emergence of life, deep subsurface microbial activity, and global geochemical carbon cycling (Ménez, 2020; Truche et al., 2020). Currently, there are two major sources of abiotic organic compounds envisioned on Earth. First, abiotic organics may be formed within the shallow crust through reduction of inorganic carbon compounds (such as carbonates or CO). Second, they could be sourced from the mantle and supplied to the crust through the migration of magmatic fluids (Sephton & Hazen, 2013). This second theory presupposes the widespread existence of carbon and organics in the mantle that were inherited from cosmic sources, such as the abiotic organic matter found in meteorites and chondrites (Gold, 1985; Sephton & Hazen, 2013). Both the first and second processes are then of interest in investigating the habitability of Enceladus, where the presumed lack of biomass means that the sole sources of organics would be either in-situ abiotic organic synthesis, or inheritance of organic material from its parent bodies.
Abiotic synthesis
Generally, the process of abiotic organic synthesis is thought to be triggered by water-rock reactions involving molecular hydrogen (H2) and a surface catalyst (McCollom, 2013; Seewald, 2001). The molecular hydrogen for these reactions may be produced abiotically by the reduction of water in the presence of Fe(II)-bearing minerals (Eq. 1.4), by reaction of FeS (pyrrhotite) with water to generate FeS2 (pyrite) plus H2, or even by the reaction of water with surface radicals during mechanical fracturing of silicate-bearing rocks (Klein et al., 2020). Water radiolysis due to radioactive decay, either by α-emitters such as U and Th or γ-ray emitter such as K40, is also considered an important source H2 in the Earth’s crust (Lin et al., 2005; Sherwood Lollar et al., 2014).
Mechanisms of abiotic organic synthesis have been and continue to be well-studied in both science and industry. Even without catalysts or mineral surfaces, equilibrium reactions between single-carbon species in water are known to form simple organic compounds of biological interest such as formic acid (HCOOH, the water-gas shift reaction in Fig. 1.7a) and formaldehyde (CH2O) (Andreani & Ménez, 2019). The presence of H2 also favors these reactions yielding reduced organic compounds (to the left in the diagram of Fig. 1.7a), although even in ideal conditions these reactions are extremely slow (Newsome, 1980; Reeves & Fiebig, 2020). The role of metal or mineral catalysts in natural systems is then the focus of most discussions on abiotic organic synthesis. For example, Fischer-Tropsch or Fischer-Tropsch type (FTT) syntheses of CH4 and straight-chain hydrocarbons (Fig. 1.7b) may occur through the following reactions:
(2n+1)H2 + nCO → CnH(2n+2) + nH2O
FTT reactions (specifically those based on the Fischer-Tropsch reaction of Eq. 1.5) are used extensively in the industrial synthesis of H2 from CO using native metal and metal alloy catalysts (Anderson et al., 1984; Schulz, 1999). This and other types of FTT reactions (including the Sabatier reaction shown in Eq. 1.6) have also traditionally been mobilized as industrial CO2 scrubbers using Fe- and Ni-catalysts (Vogt et al., 2019). Experimental and field studies have long proposed that silicate minerals may also act as a catalyst for these reactions (Hazen & Sverjensky, 2010; McCollom & Seewald, 2006; Potter et al., 2004). Field studies of both fluids at mid-oceanic ridges and off-axis hydrothermal vents (Charlou et al., 2010; Proskurowski et al., 2008) have concluded that the fluids at these vents carry abiotic CH4 and straight-chain hydrocarbons. Aside from CH4 and hydrocarbons, identifying the origin of organic species in terrestrial hydrothermal vents is complicated because of the overwhelming abundance of biotically derived organic compounds (Simoneit, 2004). Numerous studies (e.g. Aubrey et al., 2009; Marshall, 1994; Proskurowski et al., 2008) have shown the potential for amino acid synthesis at seafloor hydrothermal conditions through Strecker or Strecker-type reactions, that describe synthesis of amino acids from reaction of HCN and NH3 with a catalyst. However, unambiguous identification of any abiotically produced amino acids in seafloor hydrothermal systems remains elusive.
Inherited organics
The moons of Saturn and Jupiter are thought to be composed in part of carbonaceous chondrites, which comprise up to 2-3 % (dominantly organic) carbon constituents (Botta & Bada, 2002) and up to 1.5 % macromolecular organic matter (Chang et al., 1978; Sephton, 2002). This would provide a rich source of organic carbon to the interior of icy moons. Alteration of this organic matter in the vents of icy moons may then favor concentration and synthesis of complex compounds due to their presumably low temperatures and reducing environments, similar to the LCHF. Unlike the LCHF, the high concentration of NH3 thought to be present in Enceladus may further serve to stabilize N-bearing species, including amino acids, through inhibiting deamination reactions. For example, alanine (an amino acid) can decompose through the following deamination pathway:
Alanine (C3H7NO2) + H2O → Propionic acid (C₃H₆O₂) + NH3 Equation 1.7
From basic thermodynamic principles, increasing the NH3 on the products side of the equation will lead to a stabilization of amino acids and other amines that are present on the reactants side of the equation (Klingler et al., 2007). The low temperatures of icy satellites again favor prebiotic reactions, as decomposition reactions such as Eq. 1.7 proceed faster at higher temperatures (Abdelmoez et al., 2010; Bada et al., 1995; Sato et al., 2004). Recent models that evaluate the reaction kinetics of various amino acids in terms of the expected internal conditions in Enceladus similarly show that at low temperatures (< 25 ⁰C), the decomposition timescales for certain amino acids (proline, glutamic acid) are on the order of millions or even billions of years (Truong et al., 2019).
Table of contents :
Chapter 1. Introduction
1.1 Hydrothermal systems in the solar system and the origin of life
1.2 Models of hydrothermal activity on icy worlds
1.3 Experimental insights into olivine alteration on icy worlds
1.4 Abiotic organic synthesis and alteration
1.5 Problematic
1.6 References
Chapter 2. 1-D reactive transport modelling of olivine alteration in Enceladus
2.1 Introduction
2.2 The 1-D modelling method
2.3 Results and discussion
2.4 Conclusions
2.5 References
Chapter 3. Olivine dissolution in partially frozen solutions from -20 to 22 °C
3.1 Introduction
3.2 Methods
3.3 Results and discussion
3.4 Supplementary methods and results
3.5 References
3.6 Aqueous chemistry tables
Chapter 4. Olivine dissolution and alteration in NH3-H2O from 0-150 °C
4.1 Introduction
4.2 Methods
4.3 Results
4.4 Discussion
4.5 Conclusions
4.6 Cautionary tales
4.7 References
4.8 Aqueous chemistry tables
Chapter 5. Experimental insights into NH3-olivine-organic alteration
5.1 Introduction
5.2 Glycine alteration experiments
5.3 Synthetic chondritic organic matter alteration experiments
5.4 Initial conclusions
5.5 References
Chapter 6. Conclusions and perspectives
6.1 Conclusions
6.2 Perspectives
