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date: 25 September 2020

# Chemical Weathering on Venus

## Summary and Keywords

Chemical and phase compositions of the surface of Venus could reflect a history of gas–rock and fluid–rock interactions, recent and past climate changes, and a loss of water from the Earth’s sister planet. The concept of chemical weathering on Venus through gas–solid type reactions was established in the early 1960s after the discovery of the hot and dense CO2-rich atmosphere of the planet, inferred from Earth-based and Mariner 2 radio emission data. Initial models suggested carbonation, hydration, and oxidation of exposed igneous rocks and a control (buffering) of atmospheric gases by solid–gas type chemical equilibria in the near-surface rocks. Carbonates, phyllosilicates and Fe oxides were considered likely secondary minerals. From the late 1970s onward, measurements of trace gases in the sub-cloud atmosphere by the Pioneer Venus and Venera entry probes and by Earth-based infrared spectroscopy challenged the likelihood of hydration and carbonation. The atmospheric H2O gas content appeared to be low enough to allow the stable existence of H2O-bearing and a majority of OH-bearing minerals. The concentration of SO2 gas was too high to allow the stability of Ca-rich carbonates and silicates with respect to sulfatization to CaSO4. In the 1980s, the detection of an elevated bulk S content at the Venera and Vega landing sites suggested ongoing consumption of atmospheric SO2 to surface sulfates. The supposed composition of the near-surface atmosphere implied oxidation of ferrous minerals to Fe oxides, magnetite and hematite, consistent with the infrared reflectance of surface materials. The likelihood of sulfatization and oxidation has been illustrated in modeling experiments in simulated Venus’ conditions. The morphology of Venus’ surface suggests contact of atmospheric gases with hot surface materials of mainly basaltic composition during the several hundreds of millions years since a global volcanic/tectonic resurfacing. Some exposed materials could have reacted at higher and lower temperatures in a presence of diverse gases at different altitudinal, volcanic, impact, and atmospheric settings. On highly deformed tessera terrains, more ancient rocks of unknown composition may reflect interactions with putative water-rich atmospheres and even aqueous solutions. Geological formations rich in salt, carbonate, Fe oxide, or silica will indicate past aqueous processes. The apparent diversity of affected solids, surface temperatures, pressures, and gas/fluid compositions throughout Venus’ history implies multiple signs of chemical alterations that remain to be investigated. The current understanding of chemical weathering is limited by the uncertain composition of the deep atmosphere, by the lack of direct data on the phase and chemical composition of surface materials, and by the uncertain data on thermodynamics of minerals and their solid solutions. In preparation for further atmospheric entry probe and lander missions, rock alteration could be investigated through chemical kinetic experiments and calculations of solid-gas/fluid equilibria to constrain past and present processes.

# Introduction

The composition and evolution of atmospheres of terrestrial planets is governed by delicate balances of inputs and sinks of volatile species in forms of gases, liquids, and ions. A supply of rocks and gases from the interior, escape to space, and chemical alteration of rocks are the major factors affecting the balances. Chemical alteration of surface rocks (chemical weathering) consumes volatiles to secondary minerals and sometimes releases volatiles to the environment. On the Earth and Mars, chemically altered rocks and minerals provide the main information about past ambient environments and atmospheric compositions. Likewise, chemical and phase composition of the surface materials of Venus could constrain pathways, the degree, and conditions of current and past alteration reactions, if they are understood.

Table 1. The composition of major and trace chemically active gases in the lower atmosphere of Venus based on in situ and remote measurements.

Gas

Volume fraction

Altitude

CO2

0.965 ± 0.008

<65 km

N2

0.035 ± 0.008

<65 km

SO2

(1.5 ± 0.3) × 10-4

22-42 km

H2O

(3.0 ± 1.5) × 10-5

5-45 km

CO

(1.7 ± 0.14) × 10-5

12 km

COS

(4.4 ± 1) × 10-6

33 km

H2S

(3 ± 2) × 10-6

<20 km

HCl

(0.4 ± 0.03) × 10-6

<74 km

S1-8

2 × 10-8

<50 km

HF

(5 ± 3) × 10-9

35-70 km

Note: See Fegley (2014) and Marcq et al. (2018) for data sources. Concentrations of several gases (e.g., CO, COS) change below ~30–40 km. If atmospheric gases reach gas–gas type chemical equilibria at the conditions of modal planetary radius (740 K, 95.6 bar), volume fractions of near-surface gases are: COS, (3.5–4.9) × 10−5, S2, (3–6) × 10−7, H2S, (0.6–2.7) × 10−7, and log10fO2(bar) = −21.36 ± 0.1. The equilibrium concentrations match the measured concentrations of CO2, SO2, H2O, and CO.

Although Earth and Venus could have accreted from similar materials (Lewis & Prinn, 1984; Morbidelli et al., 2012), Venus apparently lost a major fraction of its near-surface water and developed the dense and hot $CO2$-rich atmosphere with traces of other chemically active gases (Table 1) that could react with the surface materials. Although a high D/H isotopic ratio in the atmosphere suggests the dissociation and loss of water (Donahue et al., 1982), the mass of past water, conditions and timing of the loss, and effects of water on the composition and redox state of rocks remain largely unknown. Models do not exclude climate warming on Venus from putative liquid water environments (Kasting, 1988; Way et al., 2016) toward the current conditions (740 K, 95.6 bar at the planetary modal radius of 6051.4 km). An increasing solar luminosity and an increasing mass of atmospheric $CO2$ contributed to greenhouse warming through time. Large-scale degassing due to volcanism and impacts could have been a cause of episodic greenhouse warmings. A majority of current surface materials could reflect the last global volcanic and tectonic resurfacing event 0.3–1 Ga ago (Basilevsky et al., 1997) and products of gas–solid type chemical weathering formed during and after the event. Chemical weathering could also affect more ancient materials and secondary minerals that may be exposed at highly tectonically deformed tessera terrains of unknown composition.

Table 2. Chemical composition of minerals.

Mineral’s name

Ideal chemical formula

Albite

NaAlSi3O8

Andalusite

Al2SiO5

Anhydrite

CaSO4

Anorthite

CaAl2Si2O8

Apatite

Ca5(PO4)3(Cl,F,OH)

Calcite

CaCO3

Ca-Mg-Fe pyroxene

(Ca,Mg,Fe)SiO3

Corundum

Al2O3

Diopside

CaMgSi2O6

Dolomite

CaMg(CO3)2

Enstatite

MgSiO3

Fluorapatite

Ca5(PO4)3F

Fluorite

CaF2

Forsterite

Mg2SiO4

Halite

NaCl

Hematite

Fe2O3

Kalsilite

KAlSiO4

Magnesite

MgCO3

Magnetite

Fe3O4

Marialite

Na4Al3Si9O24Cl

Mg-Fe olivine

(Mg,Fe)2SiO4

Mg-Fe pyroxene

(Mg,Fe)SiO3

Microcline

KAlSi3O8

Nepheline

(Na,K)AlSiO4

Phlogopite

KMg3(AlSi3O10)(OH,F)2

Plagioclase

CaAl2Si2O8–NaAlSi3O8

Pyrite

FeS2

Pyrrhotite

Fe1-xS

Rhodochrosite

MnCO3

Rhodonite

MnSiO3

Rutile

TiO2

Siderite

FeCO3

Sodalite

Na8Al6Si6O24Cl2

Spinel

MgAl2O4

Tephroite

Mg2SiO4

Titanite

CaTiSiO5

Tremolite

Ca2Mg5Si8O22(OH)2

Troilite

FeS

Wollastonite

CaSiO3

The knowledge about chemical weathering of Venus’ surface rocks, minerals (see Table 2 for the mineral composition) and glasses is mostly obtained from remote and in situ spacecraft data on the chemical composition and physical properties of the atmosphere and surface materials. In addition to telescopic observations, data about the atmosphere and the surface have been obtained by ten flyby spacecraft, eight orbiters, fourteen atmospheric entry probes and seven landers (Fegley, 2014; see the article “The Surface of Venus”). The ongoing investigations of Venus with Earth-based telescopes and the Akatsuki orbiter are providing new data on the atmosphere and the surface. Additional information is being gained from models of current and past processes that affect the atmospheric composition and climate. Calculations of chemical equilibria in solid–gas(fluid) type systems assess a potential feasibility of alteration reactions. Laboratory studies of solid–gas(fluid) interactions constrain pathways, mechanisms and rates of alteration. Further information from Venus orbiters, entry probes, and landers will expand our knowledge about the atmosphere and surface materials.

Knowledge about chemical alteration of surface materials obtained from observations and models will inform us about a coupled atmosphere–lithosphere evolution in the present epoch, shed light on volatile-rock reactions on ancient Venus, constrain the future atmospheric evolution on the Earth at an increasing solar luminosity, and provide insights on atmosphere–lithosphere interactions on terrestrial exoplanets.

# The Origin and Development of the Field

By the early 1950s, astronomical data suggested that Venus and Earth could be sister planets with similar sizes, densities, and positions in the solar system. These similarities implied comparable bulk compositions of the bodies and analogous endogenic and exogenic processes that affect planetary surfaces. However, in contrast to Earth, infrared spectroscopic data indicated abundant $CO2$ in a waterless venusian atmosphere (Kuiper, 1949). The dryness of the atmosphere and supposedly elevated surface temperatures due to proximity to the Sun implied a possible lack of liquid water on the surface (Urey, 1952). It became evident that an absence of liquid water should strongly affect physical and chemical weathering, erosion, and sedimentation. Harold Urey (1952), the Nobel Prize laureate in chemistry and a founder of cosmochemistry, suggested that all the $CO2$ on Venus is concentrated in the atmosphere because carbonates do not form efficiently without water. In Urey’s views, carbonate–silicate reactions, for example,

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that could play a role on the water-rich Earth, where $CO2$ is mainly stored in crustal carbonates, may not control atmospheric $CO2$ on Venus. Microwave spectroscopy of the atmosphere performed from the Earth in the late 1950s and with Mariner 2 in 1962 has been interpreted in terms of a hot (> 425 °C) high-pressure atmosphere. Carl Sagan (1962) from the University of California Berkley stated that available temperature and pressure data supported Urey’s idea about a lack of crustal carbonates formed by $CO2$–rock reactions.

Figure 1. The partial pressure of $CO2$ in the near-surface atmosphere of Venus and conditions of the calcite–quartz–wollastonite equilibrium (1). The arrows show uncertainties of p$CO2$. The dashed lines show conditions at the modal radius (6051.4 km), highest, and lowest elevations.

In contrast to Urey’s initial views, geologist Robert Mueller (1963, 1964) from the University of Chicago used insights from metamorphic petrology and suggested a control of Venus’ atmospheric gases by gas–mineral type equilibria (e.g., equation 1) established through rock-alteration reactions. Mueller (1964) proposed chemical equilibration among near-surface atmospheric gases, gases, and minerals at the surface, and within the upper crust. In his model, gases do not equilibrate above a near-surface layer and the gas abundances measured in the upper troposphere with ground telescopes represent near-surface compositions. This idea implied chemical coupling of chemically active gases ($CO2$, $CO$, $H2O$, etc.) and solids, buffering of gases by larger masses of reactive solids, and a physical-chemical co-evolution of the atmosphere and the upper crust. The concept of atmosphere–crust coupling through gas–solid equilibria gained support from the first direct temperature and pressure measurements by the Soviet Venera 4 (1967) and Venera 7 (1970) entry probes that also confirmed a dominance of $CO2$ among atmospheric gases. Within uncertainties, the measured $CO2$ pressure (p) appeared to be the same as calculated for equilibrium (1) at the surface temperature of Venus (Figure 1). Chemical kinetic data for reaction (1) implied geologically rapid equilibration (Mueller & Kridelbaugh, 1973) and supported the control of atmospheric $CO2$ by minerals.

To constrain Venus’ chemical environments, Robert Mueller (1963, 1964, 1965, 1968) and John Lewis (1968, 1970) from the Massachusetts Institute of Technology (MIT) calculated chemical equilibria of chemical reactions with chemically active gases ($CO2$, $CO$, $H2O$, $SO2$, $H2S$, $S2$, HCl, and HF) from thermodynamic data of chemical species at chosen temperatures and pressures. The atmosphere–lithosphere equilibrium model allowed prediction of secondary surface mineralogy from measured atmospheric compositions. As an example, putative mineralogy of H-bearing, Cl-bearing, and F-bearing secondary minerals (amphiboles, chlorides, fluorides, etc.) has been constrained from concentrations of $H2O$, HCl, and HF obtained from Earth-based spectroscopy of the middle atmosphere. In turn, Mueller (1965) and Lewis (1970) demonstrated that the unknown composition of the lower atmosphere could be assessed from an assumed mineralogy of reactive surface and subsurface materials. These works showed that chemical equilibrium approaches could constrain pathways of chemical weathering if fresh solids appear in the coupled gas–solid system. As an example, based on the apparent dominance of atmospheric $CO2$ over $CO$, Mueller (1963, 1964) did not exclude oxidation of ferrous ($Fe2+$-bearing) minerals to ferric oxide (hematite). The oxidation–reduction (redox) conditions of the magnetite–hematite equilibrium

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were considered as an upper limit for $O2$ fugacity (f) in the near-surface atmosphere ($~10−20$ bar at 700 K). It has been suggested that atmospheric $fO2$ is controlled by abundant Fe-bearing minerals commonly used in laboratory buffers of redox conditions at elevated temperatures (Eugster, 1957). The work of Mueller (1964) implied the concentration of atmospheric $O2$ well below the telescopic detection limit of $O2$ if the atmospheric redox state is controlled by mineral equilibria. The consideration of chemical equilibria with S-bearing species set upper limits for gases and suggested oxidation of reduced Fe sulfides (troilite and pyrrhotite), formation of pyrite, and instability of Ca silicate (wollastonite) with respect to sulfatization to Ca sulfate (anhydrite) (Mueller, 1965). Although a compete gas–solid and gas–gas equilibration did not get support from further data, the Mueller–Lewis concept of atmosphere–lithosphere coupling, usefulness of chemical equilibrium calculations, effects of altitude on mineral stability and weathering reactions, and a necessity for chemical kinetic experiments set a foundation for further studies.

Table 3. Chemical composition of Venus’ surface materials at the Venera and Vega landing sites (mass %).

Venera 13

Venera 14

Vega 2

SiO2

45.1 ± 3.0

48.7 ± 3.6

45.6 ± 3.2

Al2O3

15.8 ± 3.0

17.9 ± 2.6

16.0 ± 1.8

FeO

9.3 ± 2.2

8.8 ± 1.8

7.7 ± 1.1

MnO

0.2 ± 0.1

0.16 ± 0.08

0.14 ± 0.12

MgO

11.4 ± 6.2

8.1 ± 3.3

11.5 ± 3.7

CaO

7.1 ± 0.96

10.3 ± 1.2

7.5 ± 0.7

K2O

4.0 ± 0.63

0.2 ± 0.07

0.1 ± 0.08

TiO2

1.59 ± 0.45

1.25 ± 0.41

0.2 ± 0.1

SO3

1.62 ± 1.0

0.88 ± 0.77

4.7 ± 1.5

Cl

<0.3

<0.4

≤0.3

Total

96.1

96.3

93.4

S

0.65 ± 0.40

0.35 ± 0.31

1.9 ± 0.60

Ca

5.1 ± 0.70

7.4 ± 0.86

5.4 ± 0.50

S/Ca

0.16 ± 0.12

0.06 ± 0.06

0.44 ± 0.18

Note: S/Ca is the atomic ratio.

Source: Surkov et al. (1984, 1986)

In 1978, concentrations of chemically active atmospheric gases ($CO2$, $SO2$, $CO$, $COS$, $H2O$, and $H2S$) were measured by the Pioneer Venus main probe, Venera 11, and Venera 12 entry probes (Von Zahn et al., 1983). Although most measurements were obtained above the altitude of 15–20 km, extrapolation of the data to the surface provided information to assess the stability of surface minerals. The concentration of $SO2$ (Table 1) implied weathering reactions that trap atmospheric sulfur to secondary minerals. Equilibrium models of multicomponent rock–gas systems developed at the Vernadsky Institute in Moscow (Barsukov et al., 1982; Khodakovsky et al., 1979) used a Gibbs free energy minimization method that does not consider individual reactions and provides speciation of solid (secondary mineral assemblages) and gas phases. The equilibrium calculations that included S-bearing gases suggested the formation of secondary Ca sulfate (anhydrite) and pyrite (Barsukov et al., 1982; Khodakovsky, 1982); conclusions confirmed by similar calculations at MIT (Klose et al., 1992). The trapping of $SO2$ in secondary minerals suggested from the equilibrium calculations gained support from the elevated bulk S contents of basaltic materials sampled by the Venera 13, Venera 14 (1982), and Vega 2 (1985) landers (Surkov et al., 1984; 1986) (Table 3).

Figure 2. Surface of Venus at the landing site of Venera 13. Source: USSR Academy of Sciences/Brown University.

Although the dominance of $CO2$ in the atmosphere implied relatively oxidizing conditions ($fO2>10−27$ bar at 427 °C, Mueller, 1964), the near-surface atmosphere was uncertain until the measurements of reduced gases by the Pioneer Venus and Venera probes in 1978 and interpretations of data from the Venera landers in 1980s. The inclusion of trace gas concentrations in calculations of multicomponent gas–basalt chemical equilibria suggested an oxidizing weathering of basalt leading to formation of magnetite, anhydrite, and pyrite (Barsukov et al., 1982). Color redox indicators at Venera 13 and Venera 14 landers became dark, presumably due to reduction of white $Na2V2O7$ to dark vanadium oxides, and the evaluated $fO2<10−22$ implied the stability of magnetite and/or ferrous silicates (Florensky et al., 1983b). This upper $fO2$ limit appeared to be close to values controlled by the magnetite–hematite equilibrium (equation 2). In 1986, a red reflectance slope in the near infrared spectral range (~0.7–1.1 µm) measured at the Venera 9 and Venera 10 landing sites (Golivin et al., 1983) was interpreted in terms of a reflection of hematite at the surface temperature of Venus (Pieters et al., 1986). The analysis of Venera 13 color images (Figure 2) indicated more abundant spectrally red $Fe3+$-bearing phases in surface fines (Shkuratov et al., 1987). These data suggested an ongoing oxidizing weathering of ferrous silicates (pyroxene, olivine) and glasses of igneous mafic rocks to magnetite and hematite. Subsequent evaluations of $fO2$ from atmospheric gas abundances were consistent with the redox environments near the conditions of the hematite–magnetite equilibrium (equation 2) (Fegley et al., 1997b; Zolotov, 1996).

The detection of the high D/H ratio (~0.016) in cloud aerosols by the Pioneer Venus indicated a significant escape of H throughout history (Donahue et al., 1982). Some discussed end-member scenarios included dissociation of water sourced from an Earth-like water ocean (Kasting, 1988). These interpretations suggested hydration and then dehydration of crustal rocks in Venus’ history. A massive escape of H released through the dissociation of water implies consumption of remaining oxygen through oxidation of atmospheric gases and materials in the lithosphere. These inferences suggested oxidation of surface rocks on early Venus through gas–solid interactions and putative weathering reactions with aqueous solutions.

# Constraints from Physical Properties and Images of the Surface

The density of rocks estimated by the gamma ray densitometer at the Venera 10 landing site ($2.8±0.1gcm−3$) suggests a hard rock (Surkov et al., 1977). The density ($1.2−1.5gcm−3$), bearing strength ($2.6−10kgcm−2$), dynamic tensile strength (2.6–10 bar), and porosity (~ 50 %) of layered bedrocks estimated at the Venera 13 (Figure 2) and Venera 14 landing sites are similar to the properties of mechanically weak rocks such as volcanic tuff (Avduevsky et al., 1983; Basilevsky et al., 1985; Florensky et al., 1983a; Garvin et al., 1984; Kemurdzian et al., 1983; Surkov et al., 1984). The high porosity implies a high permeability and an enhanced surface area affected by gas–solid reactions.

The presence of rock fragments and scarce fine-grained materials seen in images (Venera 9, 10, 13, and 14 landing sites) indicates a limited physical weathering (Basilevsky et al., 1985; Florensky et al., 1977; 1983a; Garvin et al., 1984). These observations agree with the radar investigations by Venera 15 and 16 (1983–1984) and Magellan (1990–1994) orbiters that did not reveal widespread aeolian (Greeley et al., 1997; Kreslavsky & Bondarenko, 2017) and pyroclastic deposits at the ~100 m scale (Crumpler et al., 1997). However, the occurrence of parabolic impact ejecta deposits reveals a role of impacts in the production of fine-grained materials and layered deposits (Basilevsky et al., 2004). This work shows that layered rocks (Figure 2) could be lithified fine-grained materials ejected from several nearby craters and that sings of chemical weathering in surface rocks may not fully reflect alteration in situ.

Figure 3. Tessera terrain in Ovda Regio on Venus (a radar image obtained with Magellan spacecraft). The layered and deformed structures could be oceanic sediments reworked in metamorphic and/or igneous processes after cessation of an aqueous period of Venus’ history. A channel can be seen in the lower part of the image.

Source: NASA.

The mafic composition of rocks inferred by several landers (Venera 9, 10, 13, 14, and Vega 2; Surkov et al., 1984, 1986) (Table 3) is consistent with interpretations of orbital radar images. The morphology of lava flows and the lack of stratovolcanoes and Earth-like convergent boundaries of lithospheric plates with volcanic arc systems indicate a dominance of basalts on the plains, volcanic rises, and sheet volcanoes (Barsukov et al., 1986a; Head et al., 1992). Rare steep-sided domes indicated viscous lavas, possibly formed through differentiation of mafic melts toward $SiO2$-rich compositions (Parvi et al., 1992). However, the elevated concentrations of K, U, and Th at the Venera 8 landing site could signify evolved igneous rocks (Basilevsky et al., 1992; Nikolayeva et al., 1992; Shellnutt, 2019). The composition of materials that formed channels (Baker et al., 1997) and highly deformed tessera terrains (Figure 3) remains unrevealed from the radar investigations.

Further constraints on chemical weathering have been obtained from orbital radar investigations that provided information about electrical properties of surface materials (Ford & Pettengill, 1983; Pettengill et al., 1982, 1997). The dielectric constant of surface materials at the plains (δ‎ = 4.0–4.5) inferred from radar studies was consistent with that of mafic rocks without abundant secondary phases with anomalous dielectric properties such as Fe sulfides and oxides of Fe and Ti. These data suggested that abundant S-rich phases at Venera and Vega landing sites are rather sulfates than Fe sulfides and that ferrous silicates (pyroxene, olivine) and glasses are only partially oxidized to $Fe3+$-oxides. At the elevations above ~2 to 4.5 km, the observed decrease in radar emissivity (increase in reflectivity) is inconsistent with unaltered igneous rocks. The emissivity indicates a presence of phases with higher dielectric constants (δ‎ = ~20–50) (Ford & Pettengill, 1983; Pettengill et al., 1988, 1997) and/or ferroelectric minerals in which δ‎ sharply changes at a transition temperature (Shepard et al., 1994). It is unclear if this phenomenon is related to secondary minerals formed through chemical weathering (Klose et al., 1992; Pettengill et al., 1988) or condensation from the atmosphere (Brackett et al., 1995; Schaefer & Fegley, 2004).

In 1990, the Galileo spacecraft was used to observe the emission from the lower atmosphere and surface at the 1 μ‎m spectral region through several windows in the $CO2$-dominated spectrum of the planet (Carlson et al., 1993). Although the emission mainly reflects the altitude-related temperature, it could inform about physical properties of surface materials. With the Galileo data, Hashimoto et al. (2008) showed that some tessera terrains could contain higher-albedo materials compared to the basalt-dominated plains. They speculated that high-albedo materials could characterize felsic rocks formed in water-rich environments of an Earth-like continental crust. Although the thermal emission data obtained with the Venus Express orbiter (2005–2015) were consistent with Galileo data, it is unclear if the low emissivity reflects composition, grain size or roughness of surface materials (Basilevsky et al., 2012; Gilmore et al., 2015; Mueller et al., 2008). It is possible that the 1 μ‎m emissivity from highlands reflects secondary mineralogy formed through weathering.

# Current Chemical Weathering on the Plains

Solid materials exposed during the global volcanic/tectonic resurfacing event have been in contact with atmospheric gases for several hundred of Ma. The likely uniform mass of atmospheric $CO2$ favored stability of temperature–pressure conditions, though surface temperature probably increased due to a gradual rise in solar luminosity. The morphology of Venus’ surface seen in radar images indicates restricted endogenic and exogenic processes since the time of global resurfacing 0.3–1 Ga (see the article “The Surface of Venus”). Very limited masses of fresh rocky materials have been supplied, exposed, disintegrated, transported, and redeposited in the current epoch. The restricted volcanism and a suppressed volcanic degassing at the high atmospheric pressure limited disturbances in the composition of trace atmospheric gases. Cycling changes in $SO2$ concentration above Venus’ clouds since late 1970s could reflect oscillations of the atmospheric circulation (Marcq et al., 2013) rather than explosive volcanism. Abundances of reactive gases were affected by gas–solid reactions and concentrations of some species became stabilized through gas–mineral(s) type chemical equilibria. The high porosity of at least some surface materials suggests chemical weathering within a permeable layer and favors equilibration of gas–gas type reactions via mineral catalysis. Both gas–gas and gas–mineral type chemical equilibrations are more likely with increasing depth where the solid surface/gas ratio increases. An unclear but probably limited transport of elements by supercritical $CO2$ implies interactions of gases with individual mineral or glass grains through reactions controlled by diffusion of components (e.g., metal ions) in the solids. Each exposed grain contacts with an array of gases (Table 1), and the grain could be affected by both parallel and competitive gas–solid reactions. The major and assured ongoing alteration processes are oxidation of $Fe2+$ and $S2−$ in minerals and glasses, and sulfatization of Ca-bearing solids.

Table 4. Possible products of chemical weathering on Venus’ lowlands in the current epoch.

Minerals and materials

Solid products of weathering reactions

Mg-Fe olivine, (Mg,Fe)2SiO4, Mg > Fe

Forsterite, enstatite, magnetite, hematite, pyrite

Mg-Fe pyroxene, (Mg,Fe)SiO3

Enstatite, quartz, magnetite, hematite, pyrite

Ca-Mg-Fe pyroxene, (Mg,Fe,Ca)SiO3

Anhydrite, enstatite, quartz, magnetite, hematite, pyrite

Plagioclase, CaAl2Si2O8-NaAlSi3O8

Anhydrite, Na-rich plagioclase, quartz, andalusite

Basalt glass

Anhydrite, altered Ca-depleted glass, forsterite, enstatite, Na-rich plagioclase, quartz, andalusite, magnetite, hematite, pyrite

Magnetite, Fe3O4

Hematite, pyrite

Pyrrhotite, Fe1-xS

Magnetite, hematite, pyrite

Calcite, CaCO3

Anhydrite

Dolomite, CaMg(CO3)2

Anhydrite, magnesite

Siderite, FeCO3

Magnetite, hematite, pyrite

Wollastonite, CaSiO3

Anhydrite, quartz

Anorthite, CaAl2Si2O8

Anhydrite, quartz, andalusite

The majority of solids exposed to gas–solid reactions are lava flows, limited pyroclastic and impact ejecta of basaltic composition which dominate on the plains, volcanic rises and other structures. Some of these rocks could be alkali-rich mafic rocks (e.g., tephrites, Venera 13 site). Although pyroxene and plagioclase are major minerals of basalt, mafic silicate glass could be abundant due to a fast cooling of magmas in the contact with the dense atmosphere (Frenkel & Zabalueva, 1983). Less abundant felsic rocks (the Venera 8 site, steep-sided domes, tessera terrains?) would contain Na-rich plagioclase and quartz. Tessera terrains (Figure 3) may contain some sedimentary and metamorphic rocks with minerals formed before the last global resurfacing event. Bedrock on tessera could even contain carbonates, salts, H-bearing phases, $SiO2$-rich sediments, and other remnants of minerals formed in putative early wet climate. Table 4 depicts the likely alteration products of rock-forming minerals at the conditions of Venus’ lowlands.

## Oxidation Pathways

Figure 4. The stability of Fe oxides and sulfides as functions of fugacities of $SO2$ and $O2$ (bar) at the conditions of Venus’ modal radius (740 K, 95.6 bar). The arrows show the error bars due to uncertainties within the thermodynamic data. The Venus’ box corresponds to $SO2$ measurements in the lower atmosphere and to the range of atmospheric $fO2$ at the lowlands (Fegley et al., 1997b). The plot indicates the stability of the magnetite–pyrite and/or hematite–magnetite assemblages. Pyrrhotite is unstable with respect to oxidation to Fe oxides and/or pyrite.

The fugacity of $O2$ in the near-surface atmosphere has been evaluated through the consideration of gas–gas type equilibria with the use of concentrations of $CO2$, $SO2$, $CO$, $COS$, $H2S$, and $Sn$ extrapolated toward the surface, for example,

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The results of these efforts (e.g., Fegley & Treiman, 1992; Krasnopolsky & Parshev, 1981; Zolotov, 1996) summarized by Fegley et al. (1997b) led to $fO2=10−21.7$ to $10−20.0$ bar at the conditions of planetary modal radius (6051.4 km). It is unclear if the uncertainty reflects the extrapolation of measured gas concentrations toward the surface or is due to a lack of chemical equilibration between near-surface gases. Within the uncertainties of thermodynamic data, the estimated $fO2$ matches that controlled by the magnetite–hematite equilibrium (2) at the modal radius (Figure 4; Fegley et al., 1997b), consistent with the interpretation of near infrared emission at Venera 9 and Venera 10 landing sites (Pieters et al., 1986).

On the plains, magnetite and then hematite are likely oxidation products of ferrous silicates (Mg–Fe pyroxene, olivine), basaltic glasses, and reduced Fe sulfides (e.g., pyrrhotite),

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Both Fe oxides could be present among alteration products, and the magnetite–hematite assemblage may control the redox state of atmospheric gases, such as the $CO2/CO$ ratio (equation 7). Calculations of mineral stability at Venus’ plains show that Mg-rich olivine and pyroxene (e.g., $Mg1.9Fe0.1SiO4$, $Mg0.9Fe0.1SiO3$) are stable with respect to oxidation and could be oxidation products of the silicates with comparable concentrations of Fe and Mg (Klose et al., 1992; Zolotov & Volkov, 1992). Another likely oxidation process is conversion of pyrrhotite to pyrite,

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If atmospheric gases equilibrate at the lowlands, $CO2$, $H2O$ and $SO2$ have equal potential to oxidize elements. The oxidation could proceed via competitive reactions of these gases. Fegley et al. (1995) have demonstrated the feasibility of oxidation of ferrous minerals to hematite by hot $CO2$ in the first experimental studies of basalt alteration at Venus’ temperatures. Water vapor and $SO2(g)$ are also known as high-temperature oxidation agents,

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The roles of $CO2$, $H2O$, and $SO2$ in oxidation of Venus’ solids need experimental assessment.

## Sulfatization and Sulfidation Reactions

Figure 5. The stability of rock-forming silicates, carbonates and sulfates of Ca, Mg, and Na as functions of fugacities of $SO2$, $CO$, and $S2$ (bar) at the conditions of Venus’ modal radius (740 K, 95.6 bar). The equilibrium lines correspond to reactions analogous to equations (13) and (14). The arrows show the uncertainties of selected equilibria due to thermodynamic data. The Venus box corresponds to measured and estimated concentrations of gases (Table 1). Ca-bearing minerals are unstable with respect to sulfatization to $CaSO4$. Silicates and carbonates of Mg are stable. Source: Modified from Zolotov (2018).

The atmospheric pressure of $SO2$ gas is high enough to cause formation of Ca sulfate (anhydrite) through alteration of Ca-bearing minerals such as diopside, anorthite, calcite, and dolomite (Figure 5, Table 4). The ongoing sulfatization on lowlands agrees with the elevated and variable S content and S/Ca ratios measured by the Venera and Vega landers (Table 3). The diverse S/Ca ratio may reflect duration of chemical weathering, grain size and permeability of surface materials, and the phase composition of exposed rocks. Although Ca carbonates and pyroxenes are highly unstable with respect to sulfatization at all altitudes, plagioclases could be resistant, especially at the lowlands. The uncertain data on atmospheric composition (Table 1), on the thermodynamic data on anorthite, and on the plagioclase solid solution limits the evaluation of its stability. Na-rich plagioclase, Mg-rich silicates and magnesite are not affected by sulfatization, though $Fe2+$-bearing phases are subjected to oxidation or sulfidation. Mg-rich pyroxene, Na-rich plagioclase, andalusite, quartz, and reduced gases ($S2$ and/or $CO$) are expected by-products of sulfatization of basalt minerals and glasses. The following reactions illustrate possible sulfatization of diopside and calcite, if they are exposed at the surface

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The feasibility of a rapid $SO2$–calcite interaction at simulated Venus conditions was demonstrated experimentally at MIT by Bruce Fegley and Ronald Prinn (1989), in agreement with experimental data from chemical literature (Tarradellas & Bonnetain, 1973; Tullin & Ljungstroem, 1989). The likelihood of $CaSO4$ formation through $SO2$ reactions with minerals/glasses in anoxic conditions has been supported by several recent experiments (Delmelle et al., 2018; Henley et al., 2015; Palm et al., 2018; Renggli et al., 2019; Berger et al., 2019).

Figure 6. The stability of Fe-oxides and sulfides as functions of fugacities of $SO2$ and CO (bar) at the conditions of Venus’ modal radius (740 K, 95.6 bar, 0.6 km below 0 km level of 6052 km) and at the highlands (670 K, 52.5 bar, 8.5 km above 0 km level). The arrows show the error bars due to uncertainties within the thermodynamic data. The filled Venus boxes correspond to gas measurements (Table 1) and extrapolation of the CO gradient toward the surface. Plot (a) suggests the stability of the magnetite–pyrite and/or hematite–magnetite assemblages at the conditions of Venus’ plains. Pyrrhotite is unstable with respect to oxidation to Fe-oxides and/or pyrite. Plot (b) shows that pyrite is stable at the highlands and could form at the expense of Fe-oxides and pyrrhotite. Source: Modified from Zolotov (2018).

Figure 7. The stability of minerals in the Fe–O–S system as a function of altitude and volume fraction of CO (in parts per million, ppm) in the lower atmosphere of Venus. The dashed line shows the supposed gradient of CO below 12 km.

Although the low dielectric constant on lowlands precludes abundant Fe sulfides (Ford & Pettengill, 1983), secondary sulfides could contribute to the elevated S content at the Venera/Vega landing sites (Table 3). Calculations of mineral stability show that pyrite could be stable at the lowlands (Figures 4, 6, 7). Partial pressures of $SO2$ and $CO$ closely match conditions determined by the magnetite–pyrite equilibrium,

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Pyrite could form through alteration of $Fe2+$-bearing silicates, magnetite, and pyrrhotite (reactions 8–10),

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The prevalence of these and other pathways of pyrite formation is unclear. The evaluation of pyrite stability and formation/alteration reactions are limited by the uncertain concentrations of $COS$, $H2S$, and $S2$ (Table 1). Pyrite is stable if $S2$ concentration corresponds to chemical equilibrium among atmospheric gases at the lowlands (Table 1). Pyrite is not stable at the low measured concentration of $S2$ with respect to alteration to pyrrhotite (back reaction 19) (Fegley, 1997).

## Interactions with HCl and HF

Although HCl and HF are highly reactive gases, their concentrations in Venus’ atmosphere (Table 1) are not high enough to allow alteration of rock-forming silicates and carbonates of Mg, Ca, and Mn (Zolotov, 2018). The trace amounts of atmospheric HCl and HF suggest past weathering reactions that could have led to chemical equilibration of the gases with secondary halogen-bearing minerals (Fegley & Treiman, 1992; Fegley et al., 1992; Lewis, 1968, 1970; Mueller, 1968; Nozette & Lewis, 1982). Assemblages of halogen-bearing surface minerals and gas–solid weathering pathways are unknown. By analogy with terrestrial volcanic environments (Delmelle et al., 2018), past interactions of HCl and HF with silicates and glasses would lead to formation of salts ($NaCl,KCl,CaF2$ etc.) and an array of solid by-products such as quartz, andalusite, albite, corundum, and nepheline. Calculations of gas–solid chemical equilibria at Venus conditions (Fegley et al., 1992; Lewis, 1970) suggest a participation of fluorite, halite, nepheline, albite, apatite, sodalite, marialite, and F-rich amphiboles and micas. The following reactions (Mueller, 1968) illustrate these past interactions and provide examples for mineral assemblages that could regulate current abundances of involved gases,

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The formation of metamorphic halogen-bearing silicates (sodalite, scapolite, fluorphlogopite, and fluoramphiboles) (Fegley & Treiman, 1992; Lewis, 1968, 1970; Mueller, 1968) seems less likely because it requires fluidal transport of components from different altering phases. The formation of secondary apatite (e.g., Barsukov et al., 1982; Treiman et al., 2016) may not be important for a major trapping of atmospheric halogens because apatite in igneous rocks could be already rich in Cl and F.

## Formation and Weathering of Carbonates

The concentration of carbon has not been measured in the surface materials and a prevalence of carbonates can be assessed only through calculations of mineral stability and gas–solid type interaction experiments. The conditions of equilibrium (1) imply that calcite is more stable at elevation than coexisting wollastonite and quartz. If $CaSiO3$ is a component of pyroxene solid solution, calcite is stable only at the highlands. The conditions of equilibria

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allow stability of Mg-bearing carbonates only at elevation. Atmospheric $CO2$ pressure is insufficient for formation of Mn, Na or K carbonates through silicate–$CO2$ reactions. It is unclear if carbonates can form at the present surface because silicate–$CO2$ interactions are extremely slow without water fluids. Ca-bearing carbonates are subjected to rapid sulfatization (e.g., reaction 14) (Fegley & Prinn, 1989; Tarradellas & Bonnetain, 1973) and may neither form (reactions 1, 20, 21) nor survive in a permeable surface layer. The supposed lack of carbonates in weathering products of Venus’ basalts needs confirmation through investigations of the carbonate mineralogy and carbon content in situ.

The doubtful formation of carbonates at the current surface conditions does not exclude exposure of crustal carbonates in putative sedimentary, metamorphic or igneous (carbonatites) rocks. However, carbonates of Na and K are highly unstable with respect to $SO2$ (Figure 5) and HCl (Zolotov, 2018), and cannot persist at the surface. Siderite is subjected to oxidation to Fe oxides and/or sulfidation to pyrite,

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Carbonates of Mg and Mn, magnesite and rhodochrosite, are stable with respect to atmospheric $SO2$, HCl and HF, and could be present in weathering products of metamorphic rocks.

## Hydration and Dehydration Reactions

Calculations of mineral stability demonstrate that $H2O$-bearing minerals are unstable and cannot form at the conditions of Venus’ surface (Fegley & Treiman, 1992; Zolotov et al., 1997). OH-bearing end-members of mineral solid solutions (e.g., amphiboles, micas) are unstable as well. Only OH-poor mineral solid solutions could be stable, and higher OH contents are expected in minerals at the low-temperature highlands. Phlogopite (Table 2) exemplifies such solid solutions. It is unclear if complex minerals such as phlogopite can form at the surface by slow $H2O$–solid reactions that require migration of elements from different solid phases. However, some primary OH-bearing minerals in exposed igneous or metamorphic rocks could survive dehydroxylation and reach chemical equilibrium with the atmospheric $H2O$. Experiments show that even thermodynamically unstable OH-bearing minerals such as tremolite and OH-rich phlogopite could persist, owing to low rates of dehydroxylation at Venus’ conditions (Johnson & Fegley, 2000, 2003a, b). Nominally anhydrous H-bearing minerals and glasses could be present as well, though their formation and survival at the atmosphere–surface interface remains to be assessed.

# Chemical Weathering at the Highlands

Venus’ mountaintops are ~100 K colder due to the adiabatic temperature gradient and have about half the atmospheric pressure of the lowlands. The radar and 1 μ‎m emissivity at the highlands could reflect the composition of the surface and gas–solid type interactions. Highland conditions are more favorable for carbonation (reactions 1, 20, 21), sulfatization (reactions 13, 14), and hydroxylation of minerals and glasses. For example, Ca-plagioclase in basalt could be stable at the lowlands (Figure 5),

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but be subjected to sulfatization at the highlands. However, gas–solid reactions are slower at lower temperatures and gas densities, and the highlands may not be enriched in corresponding secondary phases.

Chemically active atmospheric gases (e.g., $CO2$, $SO2$, $CO$, $COS$) do not equilibrate at the highlands (Fegley et al., 1997b; Krasnopolsky, 2007; Krasnopolsky & Pollack, 1994; Zolotov, 1996) and the redox state of the atmosphere (e.g., $fO2$) may not be clearly defined. This situation could characterize lowlands as well, if gases do not equilibrate there. Surface minerals and glasses could be affected by concurrent and reverse reactions. The lack of equilibration among gases confines evaluations of redox pathways of gas–gas and gas–solid type reactions. The use of different gases in the evaluation of $fO2$ and stability of redox-sensitive minerals often leads to different results. As an example, pyrite is stable at the highlands if its stability is considered with respect to measured $SO2$, CO (Figures 6, 7), $H2S$, and $H2O$. Pyrite and/or pyrrhotite are stable if $COS$ concentration increases toward the lowlands where it equilibrates with$CO$ and $SO2$. However, pyrite is subjected to oxidation to Fe oxides if its stability is referred to measured concentrations of $COS$. In spite of these uncertainties, several gas–gas type equilibria suggest overall more reducing conditions at the highlands, which make pyrite more stable than Fe silicates and oxides (Klose et al., 1992; Zolotov & Volkov, 1992; Zolotov, 2018). Pyrite and/or pyrrhotite could form through alteration of ferrous silicates and oxides by $COS$ (equations 16, 17), $H2S$, or $S2$.

Sulfidation of Fe-bearing minerals is consistent with the radar emissivity of highlands that suggest phases with elevated dielectric constants (Ford & Pettengill, 1983; Klose et al., 1992). Another plausible explanation of the radar emissivity is low-temperature condensation of metal sulfides ($PbS$, $Bi2S3$), halides or chalcogenides from the atmosphere (Brackett et al., 1995; Schaefer & Fegley, 2004). The diversity of emissivity patterns among elevated regions may reflect variations in rock and/or atmospheric composition (Treiman et al., 2016). The emissivity data may also reflect physical processes such as accumulation of aeolian lag deposits of dense grains of Fe oxides or sulfides (Greeley et al., 1991). No consensus exists about the role of gas–solid reactions in the formation of the enigmatic mineralogy of highlands. The temperature–pressure controlled weathering and condensation reactions remain to be reconciled with the spatially variable altitude when the radar reflectivity increases, and with an abrupt decrease in the reflectivity at the highest elevations.

The nature of 1 μ‎m emissivity of elevated tessera terrains remains unclear. In addition to distinct bulk mineralogy and roughness of tessera surfaces (Basilevsky et al., 2012; Gilmore et al., 2015; Hashimoto et al., 2008), the emission could reflect a thin (≥ 1 μ‎m) coating of high-albedo minerals (albite, quartz, magnesite, etc.) formed by chemical weathering reactions. In one elevated volcanic region, the emission pattern may reflect the exposure of a geologically recent lava that has not been altered by gas–solid reactions (Smrekar et al., 2010).

# Atmosphere–Lithosphere Coupling and Evolution

## Can Surface Solids Control Atmospheric Composition?

Chemical reactions at the atmosphere–lithosphere interface imply both trapping and release of gases. The negligible atmospheric mass compared to the lithospheric mass suggests a major influence of rocks on the atmospheric composition. Reactions with solids are crucially important for chemically active atmospheric gases. These gases (except $CO2$) could be consumed in secondary minerals through gas–solid weathering reactions in a thin (< 1 m) layer of surface materials. Some gas–solid reactions could have proceeded to chemical equilibration. Any equilibration suggests coupling of atmospheric gases and surface solids, a control (buffering) of gases by solids and chemical coevolution of the atmosphere–upper crust system.

Venus’ atmosphere contains more $CO2$ than the Earth’s atmosphere, ocean and crust, and Venus’ crust may not contain carbonates that affect atmospheric $CO2$. The reactivity of Ca carbonates and silicates at the surface (reactions 13, 14) precludes control of atmospheric $CO2$ by equilibria (1 and 21) in permeable surface materials. Equilibrium (1) is unstable with respect to an increase in temperature that may cause runaway decarbonation (back reaction 1) through the greenhouse warming by the released $CO2$ (Hashimoto et al., 1997). The match of near-surface atmospheric conditions with that of equilibrium (1) (Figure 1) is probably accidental.

The match of atmospheric redox conditions (e.g., $CO/CO2$ ratio) with that of the hematite–magnetite equilibrium (equations 2 and 7) could reflect the establishment of the equilibrium through oxidation of exposed $Fe2+$-bearing phases earlier in history. At equilibrium, ratios of reduced and oxidized gases (e.g., $CO/CO2$, $H2/H2O$) remain constant if both oxides are present in surface materials and if surface temperature and pressure are constant. The $Fe2O3–Fe3O4$ equilibrium could operate in the same manner as a laboratory $fO2$ buffer (Eugster, 1957) where gas ratios do not change until consumption of a solid reactant is complete. The vast preponderance of $Fe2+$-bearing solids in Venus’ basalt implies that the atmosphere–surface system may not be oxidized beyond the conditions of the hematite–magnetite buffer.

Although S-bearing gases ($SO2$, $COS$, $H2S$, $Sn$) are among the major reactants and products of weathering reactions, it is unclear if they are fully controlled by surface mineralogy. On the one hand, the low (< 1) and variable S/Ca atomic ratio in rocks measured by the surface probes (Table 3) indicates ongoing consumption of $SO2$ (e.g., reaction 13). On the other hand, conditions of magnetite–pyrite (equation 15), hematite–magnetite–pyrite, and plagioclase–anhydrite (equation 24) equilibria are consistent with the composition of the near-surface atmosphere that assumes gas–gas type equilibration (Figures 4, 5, 6a, 7). These equilibria could have established through interaction of atmospheric $SO2$ and $COS$ with minerals and glasses.

Several data suggest a trapping of volcanic HCl and HF from the atmosphere and the establishment of gas–solid equilibria in the past. The atmospheric (Cl + F)/S ratio is much lower than in common magmas and volcanic gases. The tentative Cl content in Vega 2 solids (~0.3 wt%, Barsukov et al., 1986b) is much higher than in typical basalts. Rock-forming minerals (except Na and K carbonates) are stable with respect to low concentrations of HCl and HF in the current atmosphere (Zolotov, 2018). Atmospheric masses of HCl and HF could be consumed in a thin (< 30 cm) layer of surface materials with the concentration of < 0.1 wt% of Cl and F. It follows that current atmospheric concentrations of HCl and HF could be controlled by gas–solid equilibria with participation of $H2O(g)$, and any supply of HCl and HF from volcanic gases will be compensated by weathering reactions.

## Putative Aqueous Weathering on Early Venus

There are no solid constraints about the mass of accreted and outgassed Venus’ water, and about an occurrence of liquid water on the early planet. The following section describes possible Earth-inspired scenarios of neither vetted nor accepted geological history.

Earth could have accreted its water from a carbonaceous chondritic material also rich in C- and N-bearing organic compounds (Alexander et al., 2012). The similarity of C and N inventories on Venus and Earth suggests similar amounts of water accreted on the planets. Body dynamics models in the early solar system also suggest comparable amounts of accreted H-C-N-bearing materials on Earth and Venus (Morbidelli et al., 2000; 2012). If Venus outgassed the mass of water equivalent to the Earth’ ocean, climatic models do not exclude a prolonged existence of surface water (Kasting, 1988; Way et al., 2016). Water-rich environments could have enabled plate tectonics and the formation of $SiO2$-rich continental crust at the convergent boundaries of lithospheric plates (Campbell & Taylor, 1983). As on the Earth, aqueous processes favored formation of compositionally diverse rocks and minerals (e.g., silica, clay, carbonate or salt deposits) through uneven dissolution of protoliths, water transfer of elements, and separate precipitation of minerals with diverse solubility.

Aqueous chemical weathering on putative continents and alteration of suboceanic rocks could have produced much more secondary minerals than those predicted by gas–solid reactions. As on today’s Earth, continental weathering could be driven by meteoric, surface, and ground waters, and be facilitated by physical weathering and water erosion. One would expect weathering profiles topped by low-water solubility minerals (Al-rich clay minerals and Si, Al, Ti and Fe oxides). As on Earth, mineralogy of deeply weathered rocks may not be much different for diverse silicate protoliths. As on Earth and Mars, smectite clays could form in the medium levels of basalt weathering profiles. Erosion and surface-water streams led to accumulation of secondary phases in surface-water reservoirs. The formation of clay minerals in weathering profiles, accumulation of clay deposits, and subaqueous alteration of rocks consumed some liquid water and hydrated the crust.

An elevated surface temperature (~70–374 °C; e.g., Abe et al., 2011; Kasting, 1988) and an anoxic atmosphere determine the specifics of Venus’ weathering. $CO2$ may not dominate in the atmosphere in the presence of water (Kasting, 1988), though a low solubility in the hot ocean could account for an elevated concentration of $CO2(g)$. Both elevated ocean temperature and atmospheric $CO2$ concentration would enable acidic rainfall and formation of deeper weathering profiles similar to Earth’s laterites. Near volcanoes, S-bearing and Cl-bearing acids dominate $CO2$ in the formation of low-pH solutions. Acidic solutions are neutralized through dissolution of minerals, leading to precipitation of carbonates in the middle and lower parts of weathering crusts. Weathering reactions on continents and rock alteration in surface water reservoirs should have consumed a majority of atmospheric $CO2$ to crustal carbonates. A fraction of crustal carbonates could have returned to the atmosphere through thermal decarbonation in metamorphic and impact processes. As on Earth, the carbonate-silicate cycle (Walker et al., 1981; Berner et al., 1983) could have maintained the surface temperature through a more efficient $CO2(g)$ consumption at elevated temperature.

Without oxygenic photosynthesis and burial of organic carbon compounds, Venus’ atmosphere could have not accumulated much $O2$. However, proto-dissociation of $H2O$ in the atmosphere and H escape produced net O available for oxidation of the atmosphere and crust. The mass of accumulated O and timing of oxidation depended on the initial mass of water, H–O fractionation, and dynamics of escape. A likely major H escape before a significant increase of solar luminosity (Abe et al., 2011; Kasting, 1988) implies oxidation before ceasing of the aqueous epoch through the runaway greenhouse warming. If the atmospheric H and O is not blown away by the ultraviolet irradiation of the early Sun (Kulikov et al., 2006), partial oxidation of $Fe2+$ and sulfide S in aqueous weathering reactions on continents is a plausible scenario. Despite the elevated temperature, the oxidation could be slower than on today’s Earth owing to lower $O2$ concentrations in the atmosphere and surface waters. As in the Proterozoic eon on Earth (Lyons et al., 2014), deep parts of the ocean could have remained reduced. Whether oxidized phases accumulated in massive deposits (e.g., laterites, banded iron formations, sulfate evaporites), disseminated in the crust or subducted to the mantle depended on continental erosion and global tectonic processes.

The increasing solar luminosity and water loss through H escape, oxidation and hydration of rocks ended the aqueous period. Evaporation of surface water may have led to formation of salt deposits (evaporites) on the lowlands. Subsequent gas–solid weathering reactions reflected decreasing concentration of atmospheric $H2O$, accumulation of net O though H escape and change in the greenhouse temperatures due to changes in the atmospheric $H2O$ concentration and solar luminosity. Directions and rates of gas–solid reactions were affected by changing atmospheric pressure-temperature conditions. In turn, the weathering influenced these conditions through intake/release of gases and the greenhouse effect. Some gases could have been temporally buffered by gas–solid equilibria. Changing buffering mineral assemblage (e.g., in dehydration or decarbonation reactions) may have caused stepwise changes of the atmospheric composition and the temperature. The warming eventually led to dehydration and decarbonatization of the crust. Throughout history, both major impacts and volcanic events disrupted the atmosphere–lithosphere system and affected chemical weathering.

## Chemical Weathering in Recent and Further Geological History

The visible part of Venus’ geological history includes a geologically short time of global tectonic/volcanic resurfacing event at 0.3–1 Ga, followed by a period of diminishing endogenic activity (Basilevsky et al., 1997). During the resurfacing event, both volcanic degassing of chemically active gases and corresponding greenhouse warming of the atmosphere favored gas–solid reactions and was followed by cooling due to consumption of radiatively active gases (Bullock & Grinspoon, 1996). As on today’s Earth (Delmelle et al., 2018), films of secondary minerals formed on the surfaces of cooling lava flows and pyroclastic particles. A fraction of secondary minerals on today’s surface could have formed at the time of global basaltic volcanism and shortly after, while the atmosphere remained hot and rich in volcanic gases. In addition to contribution to the greenhouse heating, $H2O$ gas oxidized solids (reaction 11) and increased rates of other gas–solid type reactions. Coatings of Fe oxides, anhydrite, and halite could have formed on exposed basalt fragments. Some atmospheric gases (e.g., HCl and HF) could have equilibrated with secondary solids and their concentrations followed the temperature trend of global post-volcanism cooling. Concentrations of more slowly reacting gases (e.g., $SO2$) decreased gradually and may have not reached full equilibration with surface minerals. Oxidation of $Fe2+$-bearing phases in basalt by $CO2$ and $H2O$ (reactions 5–7, 11) and H escape could have led to the magnetite-hematite assemblage that probably maintains the atmospheric oxidation state at present (reaction 7).

Putative previous episodes of global volcanic activity could have been accompanied by analogues trapping of atmospheric products of volcanic degassing leading to gas–mineral equilibration for some gases. Secondary mineralogy of weathering reactions changed in time and reflected temporal physical-chemical conditions at the atmosphere–surface interface. Although tessera terrains could preserve mineralogical signs of weathering reactions that occurred before the last global resurfacing, it is unclear if any H-bearing phases could survive global warming due to volcanism, volcanic degassing and the greenhouse effect related to the global resurfacing event(s) in Venus’ history.

In the future Venus’ evolution, the coupled atmosphere–lithosphere system will be affected by episodes of local or global volcanism and greenhouse warming due to increasing solar luminosity. Changes in surface temperature will affect concentrations of gases controlled by gas–solid equilibria. As an example, atmospheric $SO2$ concentration will increase with temperature if pyrite is in equilibrium with magnetite (equation 15) and/or hematite in a permeable surface layer. The increase in $SO2$ concentration will affect chemical processes in the atmosphere and at the surface.

# Methods and Further Research

## Chemical Thermodynamic Approaches

Calculations of gas–solid type equilibria (Fegley, 2013) at Venus’ surface conditions is the common tool to assess chemical processes at the atmosphere–surface interface. The comparison of equilibrium and observed speciation assesses the deviation of natural environments from equilibrium conditions and predicts directions of chemical reactions. In one approach, the analysis of fugacity diagrams with stability fields of solid phases (Figures 1, 47) provides visualization of the comparison. Fugacity diagrams are calculated from constants of chosen gas–solid equilibria that set phase boundaries on the diagrams. Another approach is calculation of equilibrium composition of gas–solid type systems with multiple gases and solids. The equilibrium composition is calculated by the minimization of Gibbs free energy of the system that is constrained by the bulk composition of the system and thermodynamic properties of chemical compounds at chosen temperature and pressure. Volatile compounds (e.g., C, S, Cl) are set either as components of primary rocks/minerals or as fugacities of atmospheric gases. In the latter case, the system is open with respect to chosen gases and the equilibrium mineral assemblage reveals the maximum amount of atmospheric volatiles that could be trapped to secondary minerals. Both approaches can be used to analyze stability of mineral solid solutions, if activity–concentration relations in solid solutions are known from experiments and models. In all equilibrium calculations, attention needs to be devoted to uncertainties in thermodynamic values (namely Gibbs free energies of formation) which must have identical reference states and preferably be obtained from internally consistent chemical systems. Equilibrium calculations could assess directions of chemical weathering pathways in Venus’ history, including a putative aqueous era. Further applications of equilibrium calculations to Venus rely on more precise thermodynamic data on rock-forming minerals and their solid solutions (e.g., plagioclases, Fe and Fe-Ti oxides) at corresponding temperatures, and on in situ data on the near-surface compositions of gases and solids.

## Gas–Solid Reaction Experiments

Laboratory studies of gas–solid reactions further constrain directions of alteration reactions and reveal formation of intermediate and metastable products. Venus’ processes could be studied by custom-made experiments and be constrained by runs aimed to unveil industrial and terrestrial processes, such as volcanic $SO2$ reactions. Advanced experiments could provide information on formed gases, reaction rates and their temperature dependences (activation energy), reaction mechanisms, and rate-limiting stages. These studies require monitoring of gas contents and advanced methods for solid-state studies. In particular, sub-micron data on chemical profiles could reveal diffusion of elements below the reacting surface (Palm et al., 2018). There is no need to perform all modeling runs at Venus temperatures and pressures. High-temperature results could be extrapolated and low-pressure runs could be performed at Venus densities of trace gases (Fegley & Prinn, 1989). However, supercritical $CO2$–solid interactions need to be investigated at Venus conditions. In batch experiments, one problem is changes in gas composition owing to gas–gas reactions that may not equilibrate. Another problem is consumption/release of gases reacting with the sample and chamber materials. A partial solution is to use a low sample/gas ratio and less reactive (Ti, Au, Pt) materials. A high-pressure gas-flow setup that allows monitoring of gas composition provides a better solution and remains to be used for Venus studies. Interactions of individual synthetic solid phases with one or two gases will provide more fundamental data than a rock interaction with complex gas mixtures. Weathering pathways on the highlands could be constrained via experiments with disequilibrium gas mixtures. Other approaches are to monitor reacting solids in the visible and near infrared spectral ranges (Gilmore et al., 2017) or through measurements of their electrical properties by high-temperature sensors in situ. Gas-flow and inert-chamber runs may not be needed to study carbonation–decarbonation and hydration–dehydration reactions.

## Insights from Natural Analogs

Chemical weathering on Venus could be constrained thorough investigations of volcanic, metamorphic, and aqueous processes on Earth and Mars. Examples are $SO2$–solid and HCl–solid interactions in Earth’s volcanic plumes (Delmelle et al., 2018), fumaroles, and subvolcanic environments (Henley et al., 2015; Henley and Seward, 2018; McCanta et al., 2014). Data on volcanic ash suggests a rapid formation of $CaSO4$ and NaCl coatings on the surface of glassy silicate grains. Investigations of carbonate–silicate metamorphic rocks (skarns, listwanites) can shed the light on the fate of Venus’ early carbonates. Studies of contact and regional metamorphism of phyllosilicate-rich rocks provide insights on dehydration of Venus’ rocks after cessation of an aqueous period. Pre-Cambrian $SiO2$-rich seafloor deposits can constrain physical chemistry of warm oceans in the contact with $CO2$-rich and $O2$-poor early atmospheres on both planets. In particular, data on chemical and isotopic composition (e.g., S, Mo, Re) of those rocks constrain weathering conditions on continents at changing atmospheric concentrations of $O2$ and $CO2$. Geochemical studies of banded iron formations may be applicable to aqueous Venus’ history because their formation is related to low atmospheric and oceanic $O2$ levels, high $Fe2+$ concentrations in the oceans, and may be indicative of ultraviolet radiation (Bekker et al., 2014). Soda lakes in volcanic regions such as Mono Lake in California may illustrate a case for early Venus when a $CO2$-rich atmosphere affected water–rock interactions at low water/rock ratios. The recent mineralogical, chemical and isotopic studies of Mars with rovers inform about aqueous weathering in water-deficient and $O2$-deficient, $CO2$-bearing and UV-irradiated environments affected by $SO2$, $H2S$ and HCl gases released through volcanism and impacts.

## Constraints from Further Space Missions

Our understanding of chemical weathering relies on direct information from the atmosphere–surface interface. The last data from the Venus’ surface were obtained in 1985, and major results were being discussed up until the late 1990s (Fegley et al., 1997a; Wood, 1997). Limited progress occurred since then, and remaining questions cannot be solved without new data from Venus.

Figure 8. Collapsed areas on this Magellan radar image could reflect melting of subsurface chloride-rich salts (e.g., ancient evaporites) followed by formation of the outflow channels during a period of an elevated crustal and atmospheric temperature. A presence of chloride deposits in exposed rock formations will indicate past aqueous environments on Venus.

Source: NASA.

Data on the phase and elemental composition of primary and secondary minerals in rocks, rock coatings, and fine-grained materials on basaltic plains and enigmatic tesserae will largely improve our knowledge. Drilling may be needed to unveil weathering profiles in the bedrock and rock fragments. The phase composition could by assessed by X-ray diffraction and by Raman and/or reflectance spectroscopy at 0.4 to 1.2 μ‎m. Constraints will be obtained from color images and physical properties of surface materials (electrical conductivity, magnetic susceptibility, density, porosity, permeability, compressive and tensile strengths, etc.). Elemental composition can be measured by several physical methods, including X-ray fluorescent (XRF) analysis, α‎-particle X-ray spectrometry (APXS), laser-induced breakdown spectroscopy (LIBS), and γ‎-ray and neutron spectroscopy. Mössbauer spectroscopy is useful to determine the oxidation state of Fe-bearing phases. An evolved gas analysis of heated samples coupled with a mass spectrometer can provide chemical and isotopic data on S-, Cl-, C-, N- and H-bearing phases. Data on the D/H ratio in the solid samples and a detection of H-bearing phases will constrain the history of water. Data on stable isotopes of O, S, C, and Cl will further assess the fate of these elements and provenance of corresponding samples. A detection of silica-, iron oxide-, carbonate-, sulfate- or chloride-rich rocks on tessera or plains will indicate past aqueous processes and constrain paleo-environments. Eutectic melting temperature of the $NaCl–CaCl2$ salt system (777 K) is only slightly higher than the current surface temperature, and flow features (Figures 3 and 8) could be formed by melts of remobilized oceanic evaporites. Additional evidence could be obtained from concentrations and isotopic composition of redox-sensitive metals (Fe, Ni, Mn, Mo, V, U, Re etc.) in possible sedimentary rocks in tessera terrains (Figure 3). A detailed chemical and isotopic characterization of Venus’ basalts will constrain the likelihood of early Earth-like plate tectonics and assess the bulk composition of the planet.

Data on concentrations of gases in the lower atmospheric scale height (~16 km) will inform about vertical gradients, latitudinal and temporal differences, chemical equilibration among some gases, specifics of the boundary layer, and putative $CO2–N2$ mass separation in it. These data will be used to assess mineral stability and weathering pathways with chemical thermodynamic and kinetic methods. Data on the atmospheric D/H ratio and He isotopes will constrain history of water and geologically recent escape of H and He, and the data on Kr and Xe will provide constraints on an early atmospheric escape.

In addition to in situ data, new orbital spectroscopic and radar observations could assess chemical processes in the subcloud atmosphere, atmospheric circulation, and morphology, microwave and near infrared emissivity of the surface.

# Conclusions and Outstanding Questions

Figure 9. Directions of alteration of minerals in basalt at the conditions of Venus’ plains. Basalt glass could alter through oxidation, sulfatization, and sulfidation. Plagioclase with moderate and high Na content could be stable. Mg-rich pyroxene and olivine are stable with respect to chemical weathering and could be among the alteration products. Other alteration products of basalt could be presented by altered glass, quartz, and andalusite.

Data from the surface and the subcloud atmosphere of Venus suggest chemical alteration of surface rocks and minerals by atmospheric gases (Table 4, Figure 9). Ongoing sulfatization is suggested from the elevated and varied bulk S content in surface materials, calculations of mineral stability, and laboratory $SO2$–solid reaction experiments. Calcium sulfate forms in sulfatization reactions. Ongoing oxidation of ferrous minerals to iron oxides implied from calculations of gas–solid type chemical equilibria is consistent with the near-infrared reflectance of the surface measured by Venera landers. Supercritical $CO2$ could be the major oxidant at the gas–surface interface. Sulfidation of Fe-bearing phases by reduced S-bearing gases (e.g., formation of pyrite) is less certain owing to the limited data on the composition of the near-surface atmosphere. Mg-rich silicates, Na-rich plagioclases, quartz, andalusite, nepheline, albite, microcline, kalsilite, corundum, rutile, titanite, fluorite, halite, fluorapatite, anhydrite, and magnesite are stable at the surface and could be present among alteration products. In the highlands, atmospheric conditions are less suitable for oxidation but are more favorable for sulfidation of Fe-bearing phases, and sulfatization. In the lowlands, some modeled atmospheric compositions are consistent with conditions determined by the magnetite–hematite, magnetite–pyrite or magnetite–hematite–pyrite equilibrium assemblages. The agreement may not be accidental, and implies the chemical coupling of atmospheric gases with permeable surface materials that was established earlier. The low atmospheric concentrations of HCl and HF suggest a preceding trapping of the gases to minerals that may control the current gas abundances. Formation of carbonates or H-bearing phases is unlikely and atmospheric C-bearing and H-bearing gases may not be controlled by surface solids.

Table 5. Chemical weathering on Venus: Outstanding questions.

Question

What is the primary and secondary mineralogy on plains and tessera terrains?

What is the concertation of S, Cl, F, H, and C in surface materials?

What is the composition of near-surface atmosphere at different altitudes and latitudes?

Concentrations of which atmospheric gases are controlled by gas–gas and/or gas–solid equilibria?

Are microwave and 1 μ‎m emissivity patterns on the highlands related to gas–solid reactions?

How did gas–solid interactions evolve since the last global resurfacing event?

What data will be indicative of aqueous environments on early Venus?

Many of these inferences remain hypothetical and new in situ data are needed to answer outstanding questions (Table 5). Neither gas-phase nor solid-phase compositions are known at the Venus atmosphere–surface interface and new missions will need to fill this gap. In preparation for further missions, experimental studies could be aimed at understanding directions, rates and mechanisms of gas reactions with rock-forming minerals (plagioclases, pyroxenes, sulfides, etc.) and silicate glasses. Attention could be devoted to revealing the oxidizing power of $CO2$ relative to $H2O$ and $SO2$, and to the role of $CO2$ in other reactions. Insights from natural analogs remain to be made to unveil Venus’ rock alterations with and without water. New chemical equilibrium models could further assess gas–solid interactions at the present and previous epochs. Equilibrium models could constrain putative aqueous processes and subsequent gas–solid type transformations (dehydration, decarbonation, etc.) of crustal materials. These studies could be aimed at understanding mineralogical, chemical, and isotopic signs of aqueous and post-aqueous processes that could be seen in rocks, minerals, and atmospheric gases. These investigations will need to be linked to geophysical, climate, and atmospheric evolution models.

## Further Reading

Fegley, B., Jr. (2014). Venus. In H. D. Holland, & K. K. Turekian (Eds.), Treatise on geochemistry (2nd ed., Vol. 2, pp. 127–148). Amsterdam, The Netherlands: Elsevier.Find this resource:

Fegley, B., Jr., & Treiman A. H. (1992). Chemistry of atmosphere–surface interactions on Venus and Mars. In J. G. Luhmann, M. Tatrallyay, & R. O. Pepin (Eds.), Venus and Mars: Atmospheres, ionospheres, and solar wind interactions, Geophysical Monograph 66 (pp. 7–71). Washington, DC: American Geophysical Union.Find this resource:

Fegley, B., Jr., Treiman, A. H., & Sharpton, V. L. (1992). Venus surface mineralogy. Observations and theoretical constraints. Proceedings of the 22nd Lunar and Planetary Science Conference, Houston, TX, March 18–22, 1991 (pp. 3–20). (A92-30851 12-91). Houston, TX, Lunar and Planetary Institute.Find this resource:

Fegley, B., Jr., Klingelhöfer, G., Lodders, K., & Widemann, T. (1997) Geochemistry of surface–atmosphere interactions on Venus. In S. W. Bougher, D. M. Hunten, & R. J. Phillips (Eds.), Venus II: Geology, geophysics, atmosphere, and solar wind environment (pp. 591–636). Tucson: University of Arizona Press.Find this resource:

Gilmore, M., Treiman, A., Helbert, J., & Smrekar, S. (2017). Venus surface composition constrained by observation and experiment. Space Science Reviews, 212, 1511–1540.Find this resource:

Lewis, J. S., & Prinn, R. G. (1984). Planets and their atmospheres: Origin and evolution. Orlando, FL: Academic Press.Find this resource:

Marcq, E., Mills, F. P, Parkinson, C. D., & Vandaele, A. C. (2018). Composition and chemistry of the neutral atmosphere of Venus. Space Science Reviews, 214, 10.Find this resource:

Taylor, F. W. (2014). The scientific exploration of Venus. Cambridge, UK: Cambridge University Press.Find this resource:

Treiman, A. H. (2007). Geochemistry of Venus’ surface: Current limitations as future opportunities. In L. W. Esposito, E. R. Stofan, & T. E. Cravens (Eds.), Exploring Venus as a terrestrial planet (pp. 7–22). Washington, DC: American Geophysical Union.Find this resource:

Volkov, V. P., Zolotov, M. Yu., Khodakovsky, I. L. (1986). Lithospheric–atmospheric interaction on Venus. In S. K. Saxena (Ed.) Chemistry and Physics of Terrestrial Planets (pp. 136–190). New York: Springer.Find this resource:

Wood, J. A. (1997). Rock weathering on the surface of Venus. In S. W. Bougher, D. M. Hunten, & R. J. Phillips (Eds.), Venus II: Geology, geophysics, atmosphere, and solar wind environment (pp. 637–666). Tucson: University of Arizona Press.Find this resource:

Zolotov, M. Yu. (2015). Solid planet–atmosphere interactions. Treatise on Geophysics, 2nd ed., vol. 10 (pp. 411–427). Amsterdam, Netherlands: Elsevier.Find this resource:

Zolotov, M. Yu. (2018). Gas–solid interactions on Venus and other solar system bodies. Reviews in Mineralogy and Geochemistry, 84, 351–392.Find this resource:

Zolotov, M. Yu., & Volkov, V. P. (1992) Chemical processes on the planetary surface. In V. L. Barsukov, A. T., Basilevsky, V. P. Volkov, & V. N. Zharkov, (Eds.), Venus geology, geochemistry, and geophysics (pp. 177–200). Tucson: University of Arizona Press.Find this resource:

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