Sulfurbearing species play critical roles in atmospheric physicalchemical processes atmospheresurface interactions and the geological evolution of Venus. This chapter provides a comprehensive overview of 1 Instrumental data on the abundance and speciation of sulfur in atmospheric and crustal materials 2 the behavior of sulfurbearing species in the mesosphere clouds and lower atmosphere 3 chemical and mineralogical aspects of atmospheresurface interactions 4 the fate of sulfur during the formation differentiation and geological evolution of Venus including volcanic degassing gassolid reactions at the surface a putative aqueous period and subsequent evolution. The chapter also outlines outstanding questions and discusses further exploration of Venus in the context of sulfurrelevant investigations.
['Astrophysics - Earth and Planetary Astrophysics']
Sulfur on Venus Atmospheric Surface and Interior Processes
{'Mikhail Zolotov': "School of Earth and Space Exploration, Arizona State University, Tempe, Arizona 85287-1404. USA  \nE-mail: zolotov@asu.edu  \nPreprint of Chapter 16 in  \n' The Role of Sulfur in Planetary Processes: from Atmospheres to Cores '  \nedited by Daniel Harlov and Gleb Pokrovsky  \nSpringer Geochemistry  \n242 pages including 13 tables, 31 figures  \nSulfur-bearing species play critical roles in atmospheric physical-chemical processes, atmospheresurface interactions, and the geological evolution of Venus. This chapter provides a comprehensive overview of  \n- (1) Instrumental data on the abundance and speciation of sulfur in atmospheric and crustal materials,\n- (2) the behavior of sulfur-bearing species in the mesosphere, clouds, and lower atmosphere,\n- (3) chemical and mineralogical aspects of atmosphere-surface interactions,\n- (4) the fate of sulfur during the formation, differentiation, and geological evolution of Venus, including volcanic degassing, gas-solid reactions at the surface, a putative aqueous period, and subsequent evolution.  \nThe chapter also outlines outstanding questions and discusses further exploration of Venus in the context of sulfur-relevant investigations.  \n1  \n<!-- image -->", 'Outline': '2  \n<!-- image -->  \n<!-- image -->', 'Abstract': "Sulfur-bearing species play critical roles in atmospheric physical-chemical processes, atmospheresurface interactions, and the geological evolution of Venus. Most data on sulfur were obtained for the atmosphere through telescopic and spacecraft (flybys, orbiters) spectroscopic investigations and limited in situ measurements with entry probes. The atmospheric composition below ~20 km is barely known and is estimated by extrapolating measured contents toward the surface and thermochemical models. The bulk sulfur content was only measured in solid surface samples taken at three landing sites, and the geochemistry of sulfur in the interior is poorly constrained. Sulfur dioxide is the most abundant trace chemically active gas in the middle and lower atmosphere. Photochemical dissociation of CO2 above clouds to CO and O and subsequent oxidation of SO2 to H2SO4 maintains the existence of thick global clouds rich in sulfuric acid. Interaction of SO2 with basaltic glasses and Ca-bearing minerals leads to the formation of Ca and Na sulfates, consistent with the elevated sulfur content (up to ~2 wt%) in surface probes. Carbonyl sulfide, OCS, is the most abundant reduced atmospheric gas. It forms in the lower atmosphere from CO and sulfur gases (S, S2) and is oxidized by SO3, which forms via pyrolysis of H2SO4 gas below clouds. OCS could be formed and consumed via alteration and formation of sulfide minerals, respectively. In Hadley cell-type atmospheric circulation, low-latitude upwelling delivers OCS-enriched air toward clouds, and high-altitude downwelling brings OCS-depleted and CO-enriched gas to the deep atmosphere. The occurrence and fate of S1-8 gases and condensates (S8) must be more constrained. Some models suggest the formation of Sn gases below clouds, and others imply formation via photochemical reactions in upper clouds. Condensation of Sn in clouds remains to be confirmed through observations. S2 is consumed through OCS production in the deep atmosphere and released via silicate sulfatization. Although chemical disequilibria drive reactions throughout the atmosphere, some gases could equilibrate at the surface. Moreover, the abundance and speciation of S-bearing gases in the deep atmosphere could be controlled by iron sulfide-oxide and/or sulfate-silicate phase equilibria in a permeable surface layer. If this is the case, supplies and sinks of atmospheric sulfur could be compensated by buffering gas-solid reactions. Sulfidation (e.g., pyritization) and sulfatization of surface materials via gas-solid reactions are more thermodynamically favorable at highlands but need confirmation. Past volcanic degassing of SO2, OCS, S2, H2S, and/or S-Cl gases mainly constituted the current inventory of atmospheric sulfur, and the likelihood of current volcanism is questionable. Although cometary dust is a minor net source of atmospheric sulfur, it could explain elevated mixing  \n4  \n<!-- image -->  \nratios of S-bearing gases in the upper mesosphere. Cosmic sources could account for possible metal sulfates in clouds. Except for trapped atmospheric sulfur, chemical composition and radar-based morphology of immense volcanic formations suggest sulfur abundance and sulfide mineralogy (pyrrhotite, pentlandite) typical for tholeiitic and alkaline basalts. The morphology and thermal emissivity of highly tectonically deformed tessera terrains do not exclude an exposure of rocks formed in water-rich exogenic and/or endogenic environments. Whether sulfates are present in those rocks depends on water history that may or may not involve oxidizing aqueous and/or magmatic environments. Hydrogen escape from early Venus with putative water ocean could have led to sulfaterich seawater and subsequent sulfate-rich rock formations. Opposite to Earth's marine sediments, abundant pyrite might not have precipitated in Venus' counterparts due to the lack of biological sulfate reduction and ample organic matter. Examining S-bearing species and isotopes should be prioritized in future remote and in situ studies of atmospheric gases, aerosols, and chemically altered and pristine rocks. Upcoming DAVINCI, VERITAS, and EnVision missions will explore atmospheric and surface compositions, allowing a better understanding of the behavior of sulfur on the Earth's sister planet.", '16.1.1   Fates of Sulfur on Earth and Venus': "Venus and Earth have differentiated interior structures with metal-rich cores, silicate mantles, and basalt-rich crusts over most of their surfaces. These planets have secondary atmospheres and likely share significant formation steps and geological history, such as accretion from solid materials, formation of magma oceans at late stages of accretion, formation of Fe metal-rich cores, generation of secondary C-, N-rich atmospheres through magmatic degassing, and coupled physical-chemical evolution of deep interiors, silicate crusts, and atmospheres. Both planets' mantles oxidized beyond the iron-wüstite (IW) f O2 buffer conditions. On Venus and Earth, sulfur compounds are present among atmospheric gases. Sulfur is abundant in exogenic surface materials. In the atmospheres, sulfur occurs in multiple redox states, though oxidized S(VI), S(IV), and S(II) species (SO4 2, SO2, and SO) dominate. Both sulfate, S(VI) and sulfide, (S(-II) and S(-I), redox forms are most common in surface and crustal materials, though sulfides dominate deeper interiors. Exogenic S-bearing minerals are presented mainly by sulfates of rock-forming elements (Ca, Na, etc.) and pyrite. Pyrrhotite (Fe1-x) and troilite (FeS) likely compose most S-bearing compounds in crustal magmatic and mantle/core materials.  \n<!-- image -->  \nSulfur is an active player in processes that have shaped and currently affect deep interiors, crusts, surfaces, and the atmospheres. Sulfur participates in magma ocean processes; global differentiation and sequestration of Fe-rich alloys, light elements, and FeS to cores; formation of silicate mantles depleted in sulfur and chalcophile elements; separation of sulfide liquids in magmatic processes; magmatic assimilation of crustal material; volcanic degassing; trapping of volcanic gases to minerals; atmospheric thermochemical and photochemical reactions; formation of atmospheric acid sulfate aerosols; supply and transformation of space materials in the atmosphere; geochemical cycles in coupled atmosphericcrustal systems; and isotopic fractionation during these processes.  \nSulfur is an active player in processes that have shaped and currently affect deep interiors, crusts, surfaces, and the atmospheres. Sulfur participates in magma ocean processes, global differentiation and sequestration of Fe metal and FeS to cores, formation of silicate mantles depleted in sulfur and chalcophile elements, separation of sulfide liquids in magmatic processes, magmatic assimilation of crustal materials, volcanic degassing, trapping of volcanic gases to minerals, atmospheric thermochemical and photochemical reactions, formation of atmospheric acid sulfate aerosols, supply, and transformation of space materials in atmospheres, geochemical cycles in coupled atmosphericcrustal systems, and isotopic fractionation during these processes.  \nCrustal sulfur is a typical 'excess' volatile on Earth after Rubey (1951). The 'excess' means that the chemical weathering of endogenic sulfides cannot explain the mass of sulfur in the crust and the hydrosphere. This suggests a significant contribution of volcanic degassing to the inventory of sulfur in crustal materials. This concept implies the net sulfur accumulation in planetary envelopes through volcanic degassing. Although periods of global volcanic activity on Venus have ceased by ~0.3 -0.9 Ga, abundant sulfur in surface materials suggests the net accumulation of sulfur released through degassing.  \nOn both planets, sulfur abundance, speciation, and isotopic composition are indicators of physicalchemical processes throughout history and at present. These compositional data inform atmospheric chemical and photochemical processes, atmospheric circulation, redox conditions, volcanic degassing, magmatic and metamorphic processes, and chemical weathering of surface materials. Atmospheric sulfur and sulfur mineralogy of surface and crystal materials are coupled on both planets, though sulfur cycles differ.  \n<!-- image -->  \nDifferences in sizes and bulk composition, as well as in scale, specifics, and timing of geological events on Earth and Venus, are reflected in the fate of sulfur. Significant discrepancies reflect the occurrence of hydrosphere throughout most of the Earth's history, plate tectonics, organic matter in crustal materials, life, and the O2-rich atmosphere. This chapter considers the abundances of sulfur and S-bearing compounds. It discusses the fate of sulfur in Venus' atmospheric, surface, and endogenic processes at present, in geological history, and in the future. Poorly constrained processes in the interior of Venus and through the planet's history are discussed in the context of data on Earth and other solar system bodies.", '16.1.2   Venus: General Knowledge': "General knowledge of Venus can be found in books (Hunten et al. 1983; Bougher et al. 1997; Esposito et al. 2007; Taylor 2014) and review chapters (Fegley 2014; Taylor et al. 2018). Venus' distance from the Sun is ~0.7 of the Earth's distance. The orbital period is ~225 Earth days, and the planet has a slow retrograde rotation period (sideral day) of 243 days. Venus has a slightly lower modal radius (6051.4 km) and mean density (5.243 g cm -3 ) than Earth. Venus' dimensionless moment of inertia factor of 0.330.34 (Magrot et al. 2021) suggests a differentiated structure with a metal-rich core, silicate mantle, and a crust. In contrast to Earth, Venus' unimodal hypsometric curve indicates a lack of Earth-like compositional dichotomy of crustal materials with thicker felsic and thinner mafic crusts. There is no evidence of plate tectonics, and Venus likely has a stagnant- or squishy-lid lithosphere (Rolf et al. 2022). The apparent lack of intrinsic magnetic field on Venus indicates a suppressed convection in the core that could be completely solid or liquid.  \nVenus has a dense atmosphere with global sulfuric acid clouds at ~43 -73 km, depending on latitude, that limit observations of the surface from space. The mesosphere, or upper atmosphere, extends from the cloud top to ~120 km above. The presence of clouds characterizes the middle atmosphere, while the lower atmosphere spans from the cloud base down to the surface. The atmosphere mainly consists of CO2 and N2 with volume mixing ratios ( x ) of ~ 0.965 and ~ 0.035, respectively. SO2 is the most abundant S-bearing gas with a mixing ratio of ~ (1 -2) × 10 -4 , or 100 -200 parts per million by volume (ppmv). The strong greenhouse effect at higher solar luminosity accounts for a surface temperature of ~660 -760 K, depending on elevation. At the modal planetary radius of 6051.4 km, the temperature is 740 K, and the pressure is 95.6 bars.  \n7  \n<!-- image -->  \nLowlands are presented mainly by volcanic plains formed by low-viscosity lava flows. The highlands have plateaus, volcanic rises, and large shallow circular volcanic-tectonic structures (coronas) unique to Venus. About 8% of the surface consists of highly tectonically deformed plateaus called tessera. The density of about one thousand impact craters suggests a surface age of ~0.3 -0.9 Ga (Herrick et al. 2023, for review), and there is no statistically significant difference in relative ages among surface formations based on crater density. One reasonable interpretation of the crater statistics suggests geologically rapid global volcanic resurfacing. Interpretations of remote sensing data on recent and ongoing volcanism need confirmation. However, a correlation of gravity and topography data suggests active mantle plumes that contribute to the upwelling beneath significant volcanic rises. Most surface materials consist of volcanic products of mafic (basaltic) composition. The composition of tessera terrains is unknown, but they could also be composed of mafic igneous materials. Low-velocity surface winds allow only limited erosion, migration, and aeolian deposition on fine-grained materials. However, surface materials are affected by gas-solid type chemical weathering that involves trapping of atmospheric Sbearing species and reactions with C-, H-, Cl-, and F-bearing gases. Secondary sulfates, sulfides, oxides, and other compounds are suggested in altered surface materials. Low emissivity in the near-infrared (near-IR) and microwave spectral ranges implies distinct compositions of surface materials on many highlands. There is no geomorphological information on Venus' geological history before ~0.3 -0.9 Ga, and it is unclear if the planet even had a plate tectonic regime. Although the high D/H ratio in the atmosphere may indicate an early water-rich Venus, there is no solid evidence for past surface water and related processes.", '16.1.3   Data Sources and Methods to Assess Sulfur and S-bearing Species': "The need for returned samples from Venus limits the knowledge of S-bearing materials' chemical and phase composition. The data on atmospheric S-bearing gases (SO2, SO, SO3, OCS, H2S, H2SO4, Sn, CS, and CS2) and cloud aerosols (sulfuric acid and possible native sulfur and metal sulfates) have been obtained from ground-based and orbital telescopic observations in the ultraviolet (UV) to millimeter ranges, from Venus' spacecraft flyby and orbital data and entry probes, as reviewed by von Zahn et al. (1983), Esposito et al. (1983, 1997), de Bergh et al. (2006), Mills et al. (2007), Marcq et al. (2018), Titov et al. (2018), and Vandaele (2020). A significant fraction of data is obtained through remote spectroscopic observations. Although a wealth of atmospheric data is available, sulfur was only measured in three  \n<!-- image -->  \nsurface probes, and there is no apparent venusian specimen in the meteorite collection. Insights can be derived from chondritic and cometary materials, terrestrial data and processes, and experimental and numerical modeling. This section lists methods and instruments used to constrain sulfur bulk abundances and the S-bearing compounds' phase composition.", '16.1.3.1   Ground-based and Spacecraft Data on the Atmosphere': "Telescopic observations assess the speciation and concentration of S-bearing compounds from the top of the mesosphere to the lower atmosphere down to ~30 km (Table 16.1). These measurements enable monitoring atmospheric physical properties and composition and observing temporal and spatial variations between spacecraft observations. Since the initial detections of SO2 at the cloud tops with the UV observations (Barker 1979; Conway et al. 1979), multiple telescopic data revealed spatial, shortterm, and long-term concentration variations of SO2 and SO. International Ultraviolet Explorer (IUE) satellite observations (Na et al. 1990) and rocket-born UV telescopes (McClintock et al. 1994; Na et al. 1994) were used to assess SO2 and SO abundances at and above cloud top from the Earth. Hubble Space Telescope (HST) UV spectroscopy was used to constrain concentrations of SO2 and SO at ~70 -81 km (Na and Exposito 1995; Jessup et al. 2015). Telescopic IR observations at 7.4 µm, 8.6 µm, and 19 µm provided data on SO2 variability in upper clouds and cloud tops at 57 -67 km (Encrenaz et al. 2012, 2016, 2019, 2020, 2023). Telescopic near-IR observations assessed SO2 and OCS contents at cloud tops at 65 -72 km (Krasnopolsky 2008, 2010b). Microwave telescopic data constrained abundances of SO2 and SO (Sandor et al. 2010, 2012; Jessup et al. 2015; Encrenaz et al. 2015) and H2SO4 gas (Sandor et al. 2012) above clouds. Microwave radioastronomy data allowed evaluation of the concentration of H2SO4(g) down to an altitude of ~30 km (Buttler et al. 1991; Jenkins et al. 2002). Ground-based polarimetry measurements constrained the refractive index and corresponding concentration of liquid sulfuric acid in the cloud aerosol (Hansen and Hovenier 1974; Arney et al. 2014).  \nAllen and Crawford (1984) discovered narrow spectral windows in the CO2 spectrum near 1.7 and 2.3 µm in the near-IR ground-based telescopic images of Venus' night side. The finding allowed observations of SO2, OCS, and other gases (H2O, CO, HCl, and HF) below the clouds (Taylor et al. 1997). Night sky observations through near-IR spectral windows (Bézard et al. 1990, 1993; Pollack et al. 1993; Marcq et al. 2005, 2006, 2008, 2021; Arney et al. 2014) provided data on SO2 and OCS for altitudes of 52 -33 km.  \n<!-- image -->  \nOrbital observations from the Pioneer Venus 1 (Stewart et al. 1979; Esposito 1984; Esposito et al. 1979, 1988), Venus Express (Belyaev et al. 2008; Marcq et al. 2011, 2013, 2020), and Akatsuki (Yamazaki et al. 2018) in the UV spectral range provided data on spatial, temporal, and long term (since 1978) viability of concentrations of SO2 and SO at and above the cloud top ( Sect. 16.2.1.1 ). IR measurements with a Fourier spectrometer from the Venera 15 orbiter (Schaefer et al. 1990) constrained SO2 content in clouds at different latitudes (Moroz et al. 1990; Zasova et al. 1993). Solar and stellar occultation data obtained with Venus Express IR and UV spectrometers provided data on abundances, variability, and gradients of SO2, SO, and SO3 at cloud top and in the mesosphere (Belyaev et al. 2008, 2012, 2017; Mahieux et al. 2023) and on OCS upper limit at 70 -90 km (Vandaele et al. 2008). Observations during the second MESSENGER Venus flyby in the UV range were used to constrain the composition of a mysterious blue-UV absorber in upper clouds (Pérez-Hoyos et al. 2018).  \nWithin and below clouds, Venus Express orbital observations in the near-IR spectral windows constrained concentrations of SO2 and OCS down to ~30 km (e.g., Marcq et al. 2008, 2023). Akatsuki UV data were used to assess SO2 abundances and variability at the cloud top (Yamazaki et al. 2018). Gaseous H2SO4 is a good absorber of radio waves, and radio occultation data from Mariner 10 flyby in 1974 (Lipa and Tyler 1979; Kolodner and Steffes 1998) and subsequent Pioneer Venus 1 (Jenkins and Steffes 1991), Magellan (Jenkins et al. 1994; Kolodner and Steffes 1998), Venus Express (Oschlisniok et al. 2012, 2021), and Akatsuki (Imamura et al. 2017) orbiters provided data to constrain x H2SO4(g) within and below clouds down to 38 km. Venus Express polarimetry data were used to constrain the sulfuric acid concentration in the cloud aerosol (Barstow et al. 2012).  \nIn situ atmospheric measurements of S-bearing gases (SO2, OCS, H2S, Sn) were performed with the Pioneer Venus Large Probe, PVLP (1978), Venera 11 and 12 (1978), Venera 13 and 14 (1982), and Vega 1 and 2 (1985) descent spacecraft with gas chromatography (GC), mass spectrometry (MS), and gas absorption spectroscopy in the UV and visible ranges (Table 16.1). Most data were obtained for the subcloud atmosphere above 20 km. The GC measurements onboard the PVLP (Oyama et al. 1980) and Venera 12 probe (Gel ' man et al. 1980) led to the detection of SO2 at 22 -42 km. The UV absorption spectra at 220 -400 nm obtained with Vega 1 and 2 were used to estimate the SO2 profile for 0 -60 km (Bertaux et al. 1996). The GC data on H2S and OCS obtained with Venera 13 and 14 probes (Mukhin et al. 1982, 1983) exhibit inconsistencies compared to other observations. These discrepancies are often attributed to issues related to carrier gas, which compromised peak identification (Krasnopolsky 1986).  \n<!-- image -->  \nThe blue absorption (< 0.5 µm) of the solar light observed at several Venera probes in the sub-cloud atmosphere was used to estimate mixing rations of S3 and S4 in the deep sub-cloud atmosphere ( San'ko 1980; Golovin et al. 1982; Maiorov et al. 2005; Krasnopolsky 1987, 2013).  \nPVLP MS data suggested concentrated sulfuric acid as a major component of the cloud aerosol (Hoffman et al. 1980), consistent with in situ Pioneer Venus polarization and nephelometry data on the refractive index of aerosol particles (Knollenberg and Hunten 1980; Ragent and Blamont 1980). Vega reaction gas chromatography data confirmed the dominance of H2SO4 in the aerosol and were used to estimate the aerosol mass at 48 -62 km (Gel'man et al. 1986; Porshnev et al. 198 7). The X-ray fluorescent (XRF) method has been used to evaluate bulk sulfur abundance in cloud aerosols collected during descents of Venera 12 (Surkov et al. 1982a), Venera 14 (Surkov et al. 1982b), and Vega 1 and 2 (Andreichikov et al. 1987) entry probes. Mass spectrometric analysis of collected and pyrolyzed cloud particles (Surkov et al. 1986a, 1987a) on Vega 1 constrained the mass load of the sulfuric acid aerosol.", '16.1.3.2   Exploration of Surface and Crustal Materials': "The bulk sulfur content in surface materials has been measured by the XRF method at landing sites of Venera 13 (Fig. 16.1), Venera 14, and Vega 2 (Surkov et al. 1984, 1986b), though the mineralogy of any solids had not been determined. Some constraints on the original sulfur content and speciation in these samples can be inferred from the FeO/MnO ratio as an indicator of magma's f O2 (Wänke et al. 1973; Schaefer and Fegley 2017). Indirect constraints on the phase composition and abundance of sulfides in surface and near-surface materials could be obtained from the specific electrical resistance of surface materials measured at Venera 13 and 14 (Kemurdzhian et al. 1983) and Vega 2 landing sites. Dielectric properties of surface materials inferred from remote microwave observations are also informative. Microwave emissivity data of surface materials were obtained from ground-based telescopic Pettengill 1983; Garvin et al. 1985; Pettengill et al. 1982, 1988), Magellan (Pettengill et al. 1992, 1996, 1997; Klose et al. 1992; Arvidson et al. 1994; Campbell et al. 1992, 1997), Venus Express (Simpson et al. 2009), and Venera 15 and 16 orbiters. The emissivity is a function of the material's dielectric constant and may inform the composition and density of a decimeter-scale surface layer. The correlation of  \ninvestigations (e.g., Campbell et al. 1997, 1999; Pettengill et al. 1997) and the Pioneer Venus 1 (Ford and inferred dielectric properties with topography or geological features provides insights into the origins of phases with an elevated dielectric constant that could be presented by metal sulfides (Pettengill et al.  \n11  \n<!-- image -->  \n1988, 1997; Klose et al. 1992; Schaefer and Fegley 2004). Venus' emission at decimeter wavelengths measured through ground-based interferometric observations provides indirect information on dielectric properties and the composition of upper meter of surface materials, such as the occurrence of iron sulfides and oxides (Antony et al. 2022).  \nSulfur content in widespread volcanic rocks could be guessed based on magma composition suggested from the morphology of lava flows and volcanic centers seen in radar images obtained with Venera 15 and 16 (Barsukov et al. 1986a) and Magellan orbiters (Head et al. 1992; Crumpler et al. 1997; Ghail et al. 2024). Petrological types of rocks could be constrained from K, U, and Th concentrations obtained from the in situ γ -ray spectroscopy of surface materials at the landing sites of Venera 8, 9, 10, Vega 1, and Vega 2 (Surkov et al. 1987b). Rough petrological type of surface materials (e.g., mafic vs. felsic) and/or degree of their physical/chemical weathering could be estimated from nightside thermal emissivity observed through spectral windows in the CO2 spectrum at ~1 µm from Galileo (Hashimoto et al. 2008), Cassini (Baines et al. 2000), and Venus Express data (Mueller et al. 2008; Helbert et al. 2008; Basilevsky et al. 2012; Gilmore et al. 2015). Nightside surface emission data at ∼ 0.65 -0.8 μm obtained from the Parker Solar Probe Venus flybys could inform compositional differences between large regions (Wood et al. 2021; Lustig-Yaeger et al. 2023). Other assessments have been obtained from the consideration of mineral stability at the surface conditions through calculations of chemical equilibria of selected reactions and in multicomponent gas-solid type systems (e.g., Mueller 1964, 1965; Lewis 1968, 1970; Nozette and Lewis 1982; Khodakovsky 1982; Barsukov et al. 1980b, 1982c, 1986d; Fegley and Treiman 1992a, b; Klose et al. 1992; Fegley et al. 1992, 1997a; Schaefer and Fegley 2004; Zolotov 2018; Semprich et al. 2020) or experimental approaches to chemical weathering of rock and minerals by gases that demonstrate uptake of sulfur to new-formed minerals  (e.g., Fegley and Prinn 1989; Berger et al. 2019; Radoman-Shaw et al. 2022; Santos et al. 2023; Ried et al. 2024).", '16.2.1 Atmospheric Compounds': "This section reveals data on the speciation of sulfur in atmospheric gases and aerosols obtained via in situ spacecraft and remote measurements ( Sect. 16.1.3.1 ), and atmospheric processes are described in Sec. 16.3.1 . Eight S-bearing gases (SO2, SO, SO3, H2SO4, OCS, S3, S4, CS, and CS2) have been identified  \n<!-- image -->  \nbelow the altitude of 100 km (Esposito et al. 1997; Mills et al. 2007; Marcq et al. 2018; Johnson and Oliviera 2019; Vandaele 2020; Mahieux et al. 2023). The most abundant S-bearing gas is SO2 below ~70 km, OCS is the second most abundant gas below ~30 km, and H2SO4 is the second gas in clouds.  SO3, OCS, SO2, SO, CS, and CS2 are notable gases in the mesosphere (Table 16.1). Concentrations of these and other gases (Sn, H2S, S2O, ClSO2 etc.) have been evaluated through chemical kinetic (0 -112 km) and chemical equilibrium models that are best applicable for the near-surface atmosphere (e.g., Krasnopolsky and Parshev 1981a, 1983; Yung and DeMore 1982; Mills 1998; Zhang et al. 2012; Fegley et al. 1997b; Krasnopolsky and Pollack 1994; Krasnoposky 2007, 2012, 2013; Bierson and Zhang 2020). Liquid sulfuric acid, H2SO4·nH2O(l), is the main component of cloud aerosols, and other S-bearing compounds could be present (condensed sulfur, metal sulfates, and cosmogenic organic matter). Some S-bearing compounds may account for absorption in the upper clouds' blue and UV spectral ranges. In this section, speciation and concentration of S-bearing species are described in the mesosphere above cloud tops, within clouds (~47 -73 km at low latitudes and 43 -64 km at high latitudes), and in the subcloud atmosphere. Although ~75% of the atmospheric mass is located below the altitude of ~20 km, most compositional data are obtained above that altitude.", '16.2.1.1   Sulfur Oxides and Sulfur Isotopes': "Sulfur dioxide is the second chemically active gas after CO2 below ~70 km. At the cloud top and mesosphere, SO2 has been observed with the ground , Earth's orbital, and rocket telescopes and from Venus ' orbiters in the UV, IR, and microwave spectral ranges (Table 16.1). Ground-based sub-mm spectroscopy data (Sandor et al. 2010, 2012) suggest high temporal variability of SO2 at 85 -100 km with a higher x SO2 than at lower altitudes. Venus Express solar occultation data in the near-IR and UV ranges (Belyaev et al. 2012) also reveal higher mixing ratios at 85 -100 km than above clouds at 70 -85 km. Venus Express UV data on stellar occultation for 85 -105 km shows an increase in x SO2 with altitude and values 5 -4 times at the nightside than on the terminator (Belyaev et al. 2017). Microwave observations with the ALMA telescope suggest significant spatial and diurnal variations by a factor of four and a shortscale patchy pattern of x SO2 at ~88 km (Encrenaz et al. 2015). HST observations in the UV range by Jessup et al. (2015) demonstrate variable abundances of SO2 and SO at 74 -81 km. Venus Express UV data also revealed spatial and temporal SO2 variability in the mesosphere and confirmed increasing x SO2 with altitude at 70 -100 km (Vandaele et al. 2017a, 2017b). A significant increase in x SO2 from ~0.02  \n<!-- image -->  \nppmv below 90 km to ~5 ppmv at 100 km is reported from the Venus Express infrared spectrometer (SOIR) data obtained via solar occultation (Mahieux et al. 2023).  These data reveal variabilities with time and/or latitude of at least one order (typically, 2 -3) of magnitude in the mesosphere. Groundbased IR mapping of Encrenaz et al. (2012) revealed x SO2 variations within 50 -175 ppbv in the timescale of 24 h at 60 -80 km. Similar observations indicate temporal and spatial viability within 30 -700 ppmv at the cloud top (Encrenaz et al. 2016, 2019, 2020, 2023).  \nOverall, x SO2 at the cloud top (~64 -73 km) is much smaller than at the cloud base and below (Bertaux et al. 1996; Esposito et al. 1988, 1997; Marcq et al. 2018, 2020, for review). The cloud top x SO2 generally ranges from 10 to 400 ppbv, exhibiting variations across space, short-term (hourly to daily), and longterm (annual) scales, sometimes fluctuating by a factor of ten or more. Higher abundances are reported for the night side, and elevated x SO2 values are observed at the morning terminator. SO2 features associated with plumes at the cloud top have a lifetime of a few to a few tens of hours (Marcq et al. 2013; Encrenaz et al. 2016; 2019). Several observations reveal overall lower x SO2 at higher latitudes and near the sub-solar point (Esposito et al. 1979; Marcq et al. 2011, 2013, 2020, 2021; Belyaev et al. 2012; Encrenaz et al. 2019, 2023; Jessup et al. 2015) (Fig. 16.2). In contrast to other data, orbital Venera 15 IR data obtained around 1983 (Schäfer et al. 1990; Zasova et al. 1993) and sounding rocket UV observations of 1988 and 1991 (McClintock et al. 1994; Na et al. 1994) indicated elevated cloud top x SO2 at high latitudes. These observations suggest variable x SO2 typically within 10 -200 ppbv at 69 km, tens of ppbv SO2 at latitudes below ~45 °, and 100 -200 ppbv (reaching one ppmv in places) in polar regions.  \nLocal and temporal low-latitude SO2 enrichments at the cloud top are commonly associated with significantly more prominent plumes than at higher latitudes (Encrenaz et al. 2019, 2023). In addition to plumes, localized SO2 enrichments are observed above the western slopes of elevated Aphrodite Terra (Marcq et al. 2020), and concentration maps at the cloud tops indicate elevated x SO2 above other highlands (Encrenaz et al. 2023). The topography-related patterns may reflect orographic gravity waves that reach clouds ( Sect. 16.3.2.1 ).  \nGlobal cloud top x SO2 systematically observed since 1978 indicate a decline by a factor of 8 ± 4 from 1978 to 1995 (Pioneer Venus and Venera 15 orbital UV data), an increase sometime in 1995 -2006 (no observations), and a subsequent decline seen in Venus Express UV and IR data (Fig. 16.3). For 2006 -2014, a moderate decrease in averaged low latitude cloud top x SO2 values is seen in the HST UV (Jeffup  \n14  \n<!-- image -->  \net al. 2015) and IR telescopic data (Encrenaz et al. 2016), as summarized by Marcq et al. (2020). Mid-IR ground-based data of Encrenaz et al. (2020) for 64 km revealed long-term anti-correlation of SO2 and H2O cloud top abundances. However, no correlation is seen in the 2019 -2020 data (Encrenaz et al. 2023).  \nCompared to the mesosphere and cloud top, SO2 is significantly more abundant in the cloud layer (Table 16.1). For upper clouds at 62 km, Zasova et al. (1993) reported 2 -20 ppmv from Venera 15 IR data. For lower clouds at ~51 -54 km, Venus Express radio occultation data for 2006 -2014 suggest a mean x SO2 of 90 ± 60 ppmv at low latitudes and 150 ± 50 ppmv at polar latitudes (Oschlisniok et al. 2021). These data reveal a possible increase in high latitude x SO2 (and H2SO4 gas) starting with low values in 2006, a trend that disagrees with the overall decline (some spikes counteracted the downward trend, Fig. 16.3) cloud top SO2 in 2006 -2012 seen in Venus Express UV data (Marcq et al. 2013). According to Oschlisniok et al. (2021), a mean x SO2 of 180 ± 50 ppmv in lower clouds was reached in 2008 -2011 high northern latitudes, while an enhanced value of 140 ± 40 ppmv was reported for southern polar latitudes in 2009 -2012. A gradual decrease in polar x SO2 is seen after 2011/2012, with lower values in 2013. These patterns suggest long-term fluctuations in global circulation ( Sect. 16.3.1.2 ).  \nBelow the clouds, SO2 is much more abundant than in the middle atmosphere. Although SO2 variability decreases at lower altitudes, both altitudinal and latitudinal gradients are reported there. Pioneer Venus Large Probe GC data for 22 km unveil 185 ± 43 ppmv (Oyama et al. 1980), and Venera 12 GC data suggest 130 ± 35 ppmv SO2 below 42 km (Gel' man et al. 1980). These measurements are comparable with 100 -200 ppmv estimated from ground-based nightside near-IR data for 30 -52 km (Bézard et al. 1993; Polack et al. 1993; Marcq et al. 2008; Arney et al. 2014) and with analogous Venus Express data (Marcq et al. 2023). Ground-based nightside near-IR data of Marcq et al. (2021) reveal a latitudinal dependence of x SO2 at ~33 km: ~130 ppmv at 10 o S and 190 -220 ppmv at 30 -60 o N (Fig. 16.4), and the latitudinally averaged x SO2 of ~180 ppmv. Vega obtained UV spectra in situ at 10 -60 km, suggesting a strong vertical SO2 gradient (Bertaux et al. 1996). Their 120 ppmv and 150 ppmv SO2 at 43 km and 52 km, respectively, are comparable with other data (Table 16.1). However, decreasing x SO2 toward the surface to 25 -40 ppmv below 20 km inferred by Bertaux et al. (1996) is inconsistent with in situ GC data and may not be explained by atmospheric models ( Sect. 16.3.1.2 ). These in situ and remote data disagree with much higher and more uncertain x SO2 values assessed for 42 km from PVLP mass spectrometry (up to 500 ppmv, Hoffman et al. 1980) and gas chromatography (Oyama et al. 1980) that  \n<!-- image -->  \ncould characterize thermal decomposition of captured aerosols (Hoffman et al. 1980; Donahue et al. 1982; Oyama et al. 1980; Mogul et al. 2024; Zolotov et al. 2023) rather than atmospheric SO2.", 'SO': "Sulfur monoxide is detected at the cloud top and mesosphere (Table 16.1). At 64 -96 km, x SO is typically within 1 -200 ppbv. Sulfur monoxide was initially observed at the cloud top thr ough Earth's satellite UV measurements (Na et al. 1990).  Rocket UV observations of Na et al. (1994) suggested 7 -17 ppbv at 64 -96 km, which agrees with data from Na et al. (1990). Sandor et al. (2010, 2012) sub-millimeter telescopic observations revealed diurnal variations and other SO variability by an order of magnitude at 70 -100 km. These data show a decrease in the SO2/SO mixing ratio with altitude and a lower ratio on the dayside side (1.9 ± 1.2) than on the night side (37 ± 15). Venus Express UV observations of the solar occultation of Belyaev et al. (2012) indicated increasing x SO from ~1 -20 ppbv at ~85 km to ~50 -200 ppbv at 95 km and a decrease in the SO2/SO ratio with altitude. About constant x SO of 100 -200 ppbv is inferred for 95 -103 km. Further analysis of these data by Belyaev et al. (2017) suggests decreasing the SO2/SO ratio from 6 -8 at 85 km to 1 -2 at 100 km. The HST UV data for 80 km corresponds to SO2/SO of ~10 (Jessup et al. 2015). These HST coordinated with ground-based sub-mm data suggest x SO within 1 -30 ppbv at 74 -81 km. Ground-based mm-wave data of Encrenaz et al. (2015) indicated significant local and diurnal SO variations at ~88 km and suggested a uniform mixing ratio at higher altitudes. No SO is detected on the night side, consistent with its short lifetime and exclusive dayside photochemical source from SO2 dissociation ( Sect. 16.3.1.3 ). Some long-term correlation between xSO2 and the SO2/SO ratio at 70 -95 km is reported by Belyaev et al. (2017) and Jessup et al. (2015). However, the observations indicate episodic SO2 increases occurring without a simultaneous rise in SO (Jessup et al. 2015; Belyaev et al. 2017). Likewise, Encrenaz et al. (2015) reported only a partial correlation between SO and SO2 compositional maps.", 'SO3': 'SO3 is believed to be a transient gas that forms via oxidation of S-bearing gases at and above cloud tops, converts to sulfuric acid in clouds, and releases through thermal decomposition of H2SO4(g) and possibly ferric sulfates in the lower atmosphere ( Sect. 16.3.1.4 ). Based on the Venus Express near-IR data (Mahieux et al. 2023), average x SO3 increases from 0.1 ppmv at cloud tops to 10 ppmv at 95 km. No observational data on SO3 are available within and below clouds.  \n16  \n<!-- image -->', 'Sulfur Isotopes': "Sulfur isotope ratios in atmospheric gases have not been reported in the initial analysis of in situ LNMS data (Hoffman et al. 1980). Mogul et al. (2021) re-examined initial LNMS mass counts and reported 34 S/ 32 S, (5.8 ± 0.7) × 10 -2 and 33 S/ 32 S of (1.4 ± 0.9) × 10 -2 at 26 -39 km. These ratios were obtained from counts supposedly caused by the SO + ion formed in the mass spectrometer as a fragment of an Sbearing cloud aerosol material that clogged the instrument's inlet. The 34 S/ 32 S ratio corresponds to δ 34 S of 289 ± 156 % relative to the Vienna Canion Diablo troilite (V -CDT) standard value (S 34 /S 32 = 0.045) and is significantly higher than known terrestrial materials having δ 34 S of 50 % to 40 %. The substantial discrepancy with terrestrial values raises concerns about the interpretation of LNMS data in terms of isotopically heavy sulfur. Elevated counts at 49 u attributed to 33 SO could be due to H 32 SO, and some counts at 48 u could be due to H2 32 SO fragment rather than 34 SO. The 33 S/ 32 S is even less specific, and the current data analysis suggests the value of (7 ± 4) × 10 -3 (R. Mogul, private communication).  \nThe LNMS δ 34 S value of Mogul et al. (2021) is within the data range reported for cometary dust and gases (400 % to +600 %, Bergin et al. 2023) . Venus' 34 S/ 32 S ratio matches that in H2S in comet HaleBopp of (6.1 ± 0.3) × 10 -2 (Crovisier et al. 2004). The ratio also pairs with that in the atmosphere of Io, (5.95 ± 0.38) × 10 -2 , which could indicate preferential 32 S loss throughout the history of volcanic degassing of SO2, S2, and SO (De Kleer et al. 2024) and may also reflect an outer solar origin on the moon. If the interpretation of LNMS mass counts is confirmed to be accurate, it paves the way for conjecture about the contribution of cometary material to atmospheric sulfur. However, mass balance assessments indicate that space sources are not significant contributors to the atmospheric sulfur inventory ( Sect. 16.4.3 ).", '16.2.1.2   Sulfuric Acid Vapor': 'Sulfuric acid vapor is abundant within clouds and in the sub-clod layer of ~10 km thick. The abundance of H2SO4(g) is variable, typically within several ppmv, and reveals altitudinal and latitudinal patterns (Table 16.1). In clouds and down to ~30 km, H2SO4(g) is detected through ground-based microwave and sub-mm observations (Buttler et al. 1991; Jenkins et al. 2002; Sandor et al. 2012) and radio occultations obtained with Mariner 10, Pioneer Venus, Magellan, Venus Express, and Akatsuki orbiters (Jenkins and Steffer 1991; Jenkins et al. 1994; Kolodner and Steffes 1998; Oschlisniok et al. 2012, 2021; Imamura et al. 2017). The only upper limit of 3 ppbv for H2SO4(g) is obtained at 85 -100 km from ground-based  \n17  \n<!-- image -->  \nmicrowave observations (Sandor et al. 2012). Jenkins et al. (2002) used microwave observations to retrieve H2SO4(g) profiles for 0 -100 km altitudes. They reported x H2SO4(g) at ~0.1 -10 ppmv level at 30 -55 km with peak contents at 46 km and indicated a substantially higher abundance at latitudes above 45°. Mariner 10, Magellan, and Akatsuki radio occultations suggest an increasing H2SO4(g) at lower altitudes that roughly follows the saturation curve with liquid sulfuric acid in middle and lower clouds at 47 -56 km (Kolodner and Steffes 1998; Imamura et al. 2017), Fig. 16.5. Venus Express radio occultation data of Oschlisniok et al. (2012) indicates ~1 -2 ppmv H2SO4(g) within clouds (50 -55 km), and values of <1 ppmv are reported for southern polar latitudes. At higher latitudes, the highest concentrations are observed at lower altitudes, consistent with the results of Kolodner and Steffes (1998). A subsequent analysis of VeRa data by Oschlisniok et al. (2021) confirmed these findings. It showed that the topside of a sub-cloud H2SO4(g)-rich layer is located ~4 km higher at equatorial latitudes than at polar latitudes. In the equatorial region, a maximum of 12 ppmv is reported at 47 km, while polar latitudes have a maximum of 9 -12 ppmv at about 43 km. These VeRa data indicate large variability in equatorial and polar regions, while middle latitudes are characterized by lower and less variable mixing ratios (5 -7 ppmv) at 43 -47 km. These data indicate increased H2SO4(g) column density at 42 -47 km in northern polar latitudes from 2006 to 2011. No notable distinction was discovered in the global distribution of H2SO4(g) between the dayside and nightside regions in these VeRa data. For 47 -38 km, Akatsuki subcloud radio oscillations indicate an uneven decrease in x H2SO4(g) from 8 -12 ppmv at the cloud deck to 2 -5 ppmv at 38 km (Imamura et al. 2017). The behavior of H2SO4(g) gas within and below clouds is governed by saturation with respect to aerosol and thermal dissociation, respectively ( Sect. 16.3.1.5 ).', 'OCS': 'None of the in situ measurements provided reliable data on OCS. Krasnopolsky et al. (2008, 2010b) used ground-based IR spectroscopy to detect OCS at 64 -72 km and reported a highly variable content at the cloud top. Higher abundances are assessed at lower latitudes with a maximin at 0 -15° N. Venus Express near-IR, solar occultation data, initially provided the upper limit for OCS at 70 -90 km (Vandaele et al. 2008). Based on this dataset, Mahieux et al. (2023) reported an average x OCS of 1 ppbv at 65 km and a significantly higher concentration of 1 ppmv at 100 km, with a variability spanning up to four orders of magnitude. No dependence on latitude, time, or the side of the terminator has been observed, though  \n<!-- image -->  \nspatial decreases in detections were reported in the 0 -30° latitude region. The data of Mahieux et al. (2023) and Krasnopolsky et al. (2008, 2010b) are consistent in the 64 -72 km altitude range. In clouds and down to ~30 km, OCS was observed through near-IR nightside spectral windows from the ground (Bézard et al. 1990; Pollack et al. 1993; Marcq et al. 2005, 2006; Taylor et al. 1997; Arney et al. 2014) and from the Venus Express orbiter (Marcq et al. 2008, 2018, 2023). The nightside near-IR data for 33 km (Pollack et al. 1993) and 36 km (Marcq et al. 2006, Arney et al. 2014) suggest OCS gradients of -1.6 ± 0.3 ppmv/km and -0.4 ± 0.2 ppmv/km, respectively. These data indicate decreasing x OCS from ~30 to 36 km by an order of magnitude. Extrapolation of OCS gradients from (Pollack et al. 1993) and (Marcq et al. 2006) gives x OCS of 8 -36 ppmv at 29 km (Krasnopolsky 2007).  Marcq et al. (2018) discussed the applicability of a constant logarithmic gradient ( d log x OCS/ d log P ) over the 30 -37 km interval. There are indications of more abundant sub-cloud OCS at low latitudes (Marcq et al. 2005, 2006, 2008, 2023). For ~30 -40 km altitude, Marcq et al. (2023) reported higher x OCS at lower latitudes and lowest values at high North polar latitudes. No data is available for lower altitudes.', 'CO': "The mixing ratio of CO in the lower atmosphere (Table 16.1) is comparable to that of OCS at an altitude of ~30 km, though compositional gradients are opposite. The negative correlation between OCS and CO in the lower atmosphere with altitude (Pollack et al. 1993; Marcq et al. 2005, 2006, 2008; Arney et al. 2014; Marcq et al. 2018) and latitude (but not longitude, Marcq et al. 2023) indicates a significant role of CO in the atmospheric chemistry of OCS and other S-bearing gases (Yung et al. 2009; Sect. 16.3.1 ). Gas chromatography data from the PVLP (Oyama et al. 1980) and Venera 12 (Gel'man et al. 1980; Krasnopolsky 2007) suggest xCO within 12 -37 ppmv in the altitude range of 12 -42 km. These measurements indicate a decrease in xCO towards the surface, possibly reaching ~9 ppmv at modal radius conditions, as Fegley et al. (19997b) estimated. Near-IR nightside observations from the Earth (Pollack et al. 1993; Taylor et al. 1997; Marcq et al. 2005, 2006; Arney et al. 2014) and Venus Express (Marcq et al. 2008, 2018, 2023) for 33 -36 km confirmed these in situ data and the increase in x CO with altitude. Galileo, ground-based, and Venus Express near-IR observations indicate higher x CO at high latitudes for 33 -36 km (Collard et al. 1993; Arney et al. 2014; Marcq et al. 2005, 2006, 2008, 2018, 2021, 2023), Fig. 16.4. Based on the comprehensive data from Venus Express, a notable difference of ~15 ppmv has been assessed between the equatorial and polar regions (Tsang and McGouldrick 2017). Interestingly, observations of Marcq et al. (2023) for lower latitudes suggest a zonal variability CO (and  \n19  \n<!-- image -->  \npossibly OCS) correlated with the longitudinally shifted average surface elevation, yielding an effect of topography on circulation. Models suggest CO formation via CO2 photolysis above ~70 km and oxidation of OCS below clouds and consumption through conversion to OCS closer to the surface ( Sect. 16.3.1 ).", 'Sn gases': "The strong absorption observed in low-resolution spectra at 0.45 -0.6 µm obtained in situ with Venera 11, 12, 13, and 14 probes below ~30 km have been interpreted in terms of light absorption by S3 and S4 gases (Moroz et al. 1981, 1983 ; Golovin et al. 1982; San'ko 198 1; Maiorov et al. 2005; Krasnopolsky 1987, 2013). Moroz et al. (1981) evaluated xS3 at 3 -15 km (Table 16.1). The interpretation of absorption spectra in terms of concertation required knowledge of the speciation of sulfur at Venus' conditions. Calculations of chemical equilibria in the Sn system by San'ko (1981) based on thermodynamic data of Mills (1974) suggested xS3/xS4 of ~ 10 4 at 0 -40 km and implied a significant contribution of S3 to absorption. Zolotov (1985) calculated the speciation of the Sn (n = 1 -8) system based on Gurvich et al. (1989 -1994) and found much lower x S3/ x S4 ratios and comparable concentrations of these gases above ~25 km. Krasnopolsky (1987) assessed an absorption cross-section of S3 to constrain its abundance at 5 -25 km by considering Sn speciation based on Mills (1974). Maiorov et al. (2005) separated the effect of Rayleigh scattering from true sulfur absorption and inferred x S3 and x S4 based on Venera 11 spectra. Krasnopolsky (2013) noted that using the thermodynamic data of Mills (1974) results in extremely low S4 densities argued for a higher x S4/ x S3 ratio at chemical equilibrium. He used a new experimental absorption cross-section of Sn gases and sulfur absorption spectra from Maiorov et al. (2005) to evaluate x S3 and x S4 at 3 -19 km. All these evaluations suggest x S3 and x S4 within 10 -10 -10 -12 .  \nAlthough S2 gas is not detected anywhere in the atmosphere, it is the most abundant Sn gas in the nearsurface atmosphere according to chemical equilibrium calculations (Table 16.2; San'ko 198 1; Zolotov 1985). If x S2 is calculated from xS3, x S2 = (0.01 -0.02) ppmv (Golovin et al. 1982; Moroz et al. 1981, 1983; Zolotov 1985; Model 10 in Table 16.2); if it is calculated from measured SO2, CO, and CO2 abundances, x S2 = (0.1 -0.4) ppmv.", 'H2S': "The concentration of H2S is insufficiently constrained from remote and in situ measurements. Venus Express solar occultation data provide highly uncertain upper limits at 60 -100 km (Mahieux et al. 2023), and Krasnopolsky (2008) assessed an upper limit for cloud top based on ground near-IR observations  \n<!-- image -->  \n(Table 16.1). The interpretation of the Pioneer Venus Large Probe LNMS data initially suggested an H2S (if counts at 34 u correspond to H2S) gradient in the sub-cloud atmosphere, an average of 3 ± 2 ppmv from 24 km to the surface, and about one-third of this value in clouds (Hoffman et al. 1980).  The subsequent LNMS report mentions only a marginal detection of ~1 ppmv H2S in clouds (von Zahn et al. 1983). PVLP GC data provided an upper limit of 2 ppmv at 22 km (Oyama et al. 1980), and H2S is absent from Ve nera 12 CG data (Gel'man et al. 1980). The reported 80 ± 40 ppmv H 2S together with abundant O2 and H2 based on in situ Venera 13 and Venera 14 GC data at 29 -37 km (Mukhin et al. 1982, 1983) are commonly considered incorrect (Krasnopolsky 1986) and have not been mentioned in comprehensive reviews of Esposito et al. (1997), Mills et al. (2007), and Marcq et al. (2018).  All reported xH2S values at and above ppmv level have not been independently confirmed and are inconsistent with atmospheric models ( Sect. 16.3.1.9 ). Mogul et al. (2021) reported possible signs of H2S in the LNMS mass counts at 34 u related to middle cloud samples, especially if PH3 (phosphine) is absent. As with SO2, the suspected presence of H2S in gases analyzed below clouds could reflect the thermal decomposition of captured non-H2SO4 cloud aerosols rather than atmospheric composition (Zolotov et al. 2023). In addition to H2S, the interpretation by Mogul et al. (2021) does not exclude a variety of reduced gases (e.g., HCN, C2H4, NH3, PH3) that are not expected to be abundant in the clouds but could be released from heated aerosols. All in situ data on H2S may reflect the thermal decomposition of captured aerosols, consistent with Vega GC data (Petryanov et al. 1987) on H2S release from heated aerosols ( Sect. 16.2.1.4 ). Chemical equilibrium models for the near-surface atmosphere (Table 16.2) suggest an order of magnitude less x H2S than reported by Hoffman et al. (1980).", '16.2.1.4   Aerosols in Clouds and Hazes': 'Multiple remote and mission data provided information on the physics of clouds and hazes (refractive index, particle size, mass loads, layering, plumes, etc.), as reviewed by von Zahn et al. (1983), Esposito et al. (1983, 1997), Krasnopolsky (1989, 2006), and Titov et al. (2018). In the 1970s, information on the cloud vertical profile was obtained via nephelometry abroad Venera 9, 10, and 11 (Marov et al. 1980) and the four Pioneer Venus entry probes (Ragent and Blamont 1980), and from the PVPL cloud particle size spectrometer (Knollenberg and Hunten 1979). The analysis of the Pioneer Venus data by Knollenberg and Hunten (1980) revealed a 20 km thick primary cloud (47.5 -70 km), surrounded by lower density upper (70 -90 km) and lower hazes (31 -47.5 km) (Fig. 16.6). According to their data, the  \n21  \n<!-- image -->  \nprimary cloud consists of upper (56.5 -70 km), middle (50.5 -56.5 km), and lower cloud (47.5 -50.5) layers. Particles have distinct sizes that may reflect different compositions, 0.3 -0.4 µm (Mode 1), 2 -2.5 µm (Mode 2), and 7 -8 µm particles that contain the bulk of cloud mass loading (Mode 3).  All size Modes are suggested in the middle and lower clouds, Modes 1 and 2 are inferred in the upper clouds, and Mode 1 particles are the presumed constituents of the upper and lower hazes.  Based on optical properties, Knollenberg and Hunten (1980) noted that Mode 1 particles could be rich in S-rich compounds, Mode 2 likely consists of spherical droplets of sulfuric acid, and Mode 3 could be presented by non-spherical solid grains.  \nThe cloud structure obtained with the particle size spectrometer and nephelometer onboard Vega probes (Moshkin et al. 1986; Gnedykh et al. 1987) were generally consistent with the Pioneer Venus data and confirmed a sharp boundary of primary cloud and the presence of hazes down to ~30 km. Though only two particle size modes (0.25 -2.5 µm and 1 -2.5 µm) were reported in the middle and lower cloud, the Mode 2 particles were significantly less numerous than in Pioneer Venus data, and nonsphericity of small particles was suggested by Gnedykh et al. (1987). Vega data implies a dense (0.1 -2 mg m -3 ) sub-cloud haze of small (< 0.25 µm) particles down to ~35 km. The increasing aerosol mass load from the upper to lower cloud regions was inferred through a collection of aerosols on filters onboard Venera 12 (Surkov et al. 1982a) and Vega (Andreichikov et al. 1987) probes. Vega in situ measurements of UV light absorption from 63 km to the surface reported by Bertaux et al. (1996) confirmed the layered cloud structure, consistent with the Pioneer Venus data (Fig. 16.6). Other investigations indicated latitudinal dependence of cloud structure and altitudinal boundaries (e.g., Cottini et al. 2012). Venera 15 IR observations suggest Mode 3 particles only at low altitudes (Zasova et al. 2007). All listed in situ data on cloud physics imply spatial and temporal variability. In addition to hazes above and below clouds, Grieger et al. (2004) suspected a near-surface haze layer at 1 -2 km altitude based on Venera 13 and 14 in situ spectrophotometry.', 'Bulk Sulfur Content, other Elements, and Aerosol Decomposition Products': "The presence of condensed S-bearing compounds in clouds is supported by the detection of bulk sulfur by the XRF method in aerosols collected on filters by Venera 12, Venera 14, Vega 1, and Vega 2. The interpretation of Venera 12 data suggested the presence of chlorine (Petryanov et al.  1981a) and iron (Petryanov et al.  1981b). Based on this experiment, Surkov et al. (1982a) did not report any iron but estimated a sulfur mass load of 0.1 mg m -3 and a Cl/S mass ratio of ~21 in aerosol collected at 47 -54 km.  \n<!-- image -->  \nIn a comparative study, Surkov et al. (1982b) estimated sulfur mass loads to be 1.1 ± 0.13 mg m -3 and determined the Cl/S mass ratio of ~ 0.09 in Venera 14 samples collected at 47 -63 km. These data indicated the highest aerosol density at 47 -56 km. Vega XRF data reported by Andreichikov et al. (1987) suggested abundant sulfur in the primary cloud aerosol at 47 -61 km, while the presence of Cl, Fe, and P was suggested. Their data analysis implied an alternating dominance of S, Cl, and P in narrow cloud layers, though the validity of such model-based interpretation was questioned by Krasnopolsky (1989). The dominance of phosphorus over sulfur at 47 -51 km implied by Andreichikov et al. (1987) and Venera 12 data on the high Cl/S ratio (Surkov et al. 1982a) disagrees with x H2SO4(g) values ( Sect. 16.2.1.2 ) at equilibrium with sulfuric acid in the middle and lower cloud regions (Fig. 16.5) and with other measurements described below.  \nThe presence of condensed S-bearing compounds in primary cloud aerosol has been suggested from in situ mass spectrometry and gas chromatography data at PVLP and Vega probes. Although the LNMS instrument has not been designed to analyze aerosols, aerosols entered the inlet in the primary cloud and clogged the inlet at ~51 km (Hoffman et al. 1980; Mogul et al. 2023, 2024). These interpretations suggest that S-bearing fragments identified in the mass counts below ~51 km represent thermally decomposed aerosols captured and temporarily stored in the LNMS inlet assembly. The mass counts suggest the release of SO2, H2O, and trace S-bearing gases (e.g., H2S, Sn) from the aerosols (Hoffman et al. 1980; Donahue et al. 1982; Mogul et al. 2021, 2024; Zolotov et al. 2023).  Significantly higher than the atmospheric concentrations of SO2 and H2O reported for the 42 km sample by PVPL GC measurements could also reflect the decomposition of aerosol captured in the primary cloud (Oyama et al. 1980; Mogul et al. 2024).  Vega 1 mass spectrometric analysis of collected and pyrolyzed (up to 673 K) aerosols suggested the release of S-O-bearing and Cl-bearing gases, and bulk 2 -10 mg m -3 S and > 0.3 mg m -3 Cl were estimated for the lower cloud aerosol (Surkov et al. 1986a, 1987a). In another investigation, Vega 1 and 2 aerosols were collected in two separate cells, heated to 353 K and 573 K, and the released water vapor and SO2 were detected in Vega 2 samples with gas chromatography (Gel'man et al. 1986). In addition to H2O and SO2, Porshnev et al. (1987) reported traces of H2S and OCS (H2S/SO2 = OCS/SO2 of ~10 -3 ) released from heated Vega 2 aerosols.", 'Sulfuric Acid': "The presence of liquid sulfuric acid as a major cloud aerosol component was inferred from ground-based polarimetry and near-IR spectroscopy, the refractive index of clouds, mass spectrometry, and gas  \n<!-- image -->  \nchromatography data on captured aerosols. Observations commonly suggest >75% sulfuric acid (H2SO4 mass % in H2SO4·nH2O) and altitudinal, longitudinal, spatial, and temporal variations in the composition.  \nYoung (1973) first noted that the IR spectra of Venus are consistent with the presence of sulfuric acid in upper clouds. Hansen and Hovenier (1974) interpreted ground-based polarimetric data at visible and near-IR wavelengths regarding scattering from 1 µm particles with a refractive index of 75% sulfuric acid. The spectroscopic ground-based near-IR data of Pollack et al. (1978) and Krasnopolsky (2008) inferred a refractive index of 1 µm practices corresponding to 85% H2SO4 at cloud tops. The nightside Venus Express near-IR observations suggested 75 -85% acid in optically thin upper clouds and 90 -100 wt% in deeper optically thick clouds (Barstow et al. 2012). These observations also indicate increasing the concentration from 80 wt% at low latitudes to 90 wt% poleward of 60° S. The interpretation of this dataset by Cottini et al. (2012) for ~1 µm cloud top droplets suggested acid in the interval 75 -85% and a narrow range of 80 -83% at low latitudes (±40°). The analysis of these data by McGouldrick et al. (2021) suggests >60% acid in upper clouds and 85 -90% in lower clouds, and a latitudinal dependence is observed. The nightside ground-based near-IR observations of Arney et al. (2014) were used to map H2SO4 concentration in the aerosol and reported an average sulfuric acid concentration of 79 ± 4%. These observations indicate a correlation of H2SO4 concentration with highly variable cloud opacity and reveal slightly more systematic changes with altitude and latitude.  More concentrated acid is tentatively suggested in the northern hemisphere and at low latitudes. A positive correlation of SO2 gas content with altitude revealed from Venera 15 orbital IR observations of Zasova et al. (1993) is consistent with the sulfuric acid aerosol (75 -85%) as a significant component of upper clouds (Mills et al. 2007). An 85% sulfuric acid in lower clouds was inferred from the amount of SO2 and H2O released from the clogged LNMS inlet assembly at the Pioneer Venus Large Probe (Hoffman et al. 1980).  \nResults of Vega 1 reaction gas chromatography analysis confirmed the dominance of H2SO4 in collected and heated aerosols with an averaged H2SO4 mass load of 0.6 ± 0.1 mg m -3 at 48 -61.5 km, though significantly lesser loads were inferred from Vega 2 reaction GC data (Gel'man et al. 1986; Porshnev et al. 1987). The mass spectrometric detection of abundant SO2 (64 u) released from pyrolyzed Vega 1 aerosol agreed with the sulfuric acid composition and suggested a lower limit of aerosol H2SO4 mass load of 2 mg m -3 (Surkov et al. 1986a, 1987a). The presence of sulfuric acid in the primary cloud aerosols agrees with physical chemistry of the H2SO4-H2O system that is consistent with measured x H2SO4(g) ( Sect. 16.2.1.2 , Fig. 16.5) and with modeled vaporization of liquid aerosol particles at temperature-  \n24  \n<!-- image -->  \npressure conditions at ~48 km (Krasnopolsky and Pollack 1994; Krasnopolsky 2015; Dai et al. 2022) ( Sect. 16.3.1.5 ). The composition of hazes above and below clouds is not known, though metastable H2SO4 aerosols are suspected to be the main component of upper hazes at 70 -90 km (Luginin et al. 2016). However, elevated concentrations of SO3, SO, SO2, OCS, CS, and CS2 in the upper mesosphere (Table 16.1) are more consistent with cosmogenic origins of mesospheric S-bearing aerosols.", 'Other S-bearing Species': 'Although only sulfuric acid is detected in the clouds, several physical and compositional data suspect other S-bearing species in aerosols, especially in Mode 1 and Mode 3 particles. A long list of S-bearing compounds has been considered to account for a blue-UV absorber in upper clouds. However, no candidate species have been detected yet ( Sect. 16.3.1.8 ). Condensed native sulfur (S 0 ), often called polyatomic sulfur or polysulfur (S8, Sx), is commonly considered a vital cloud aerosol constituent in chemical, photochemical, and physical models ( Sect. 16.3.1.7 ). Condensed polysulfur in clouds was suggested based on reaction gas chromatography analysis of captured aerosols on the Vega 1 and 2 probes that indicated the H2SO4/Sn mass ratio of ~7 -10 in aerosol (Porshnev et al. 1987). Gnedykh et al. (1987) suggested the sulfuric acid to S 0 mass ratio of ~10 from Vega nephelometer and particle size spectrometer data. Knollenberg and Hunten (1980) mentioned crystalline sulfates as candidates for Mode 3 grains. In addition to possible Fe-, Cl-, and P-containing compounds discussed above, non-acid Sbearing aerosol constituents were suspected based on PVPL mass spectrometry (Hoffman et al. 1980; Mogul et al. 2021, 2024) and Vega 2 gas reaction gas chromatography (Gelman et al. 1986; Porshnev et al. 1987). The release of H2S and OCS from captured aerosols (Hoffman et al. 1980; Porshnev et al. 1987) suggests inorganic and/or organic compounds with reduced sulfur (S 0 and/or S 2). The apparent presence of abundant S-bearing species (SO2, S, H2S, SO3, etc.) suspected in the LNMS mass counts below ~25 km could be products of thermal decomposition of non-H2SO4 aerosols (e.g., ferric sulfates, cosmogenic organic matter) captured from clouds and sub-cloud hazes (Mogul et al. 2024; Zolotov et al. 2023).', '16.2.2 Surface and Crustal Materials': 'The radar-inferred geomorphology shows the dominance of basaltic materials on plains and other volcanic formations. Venera/Vega lander data suggest sulfur abundance and mineralogy typical for planetary basalts with a contribution of exogenic sulfur. The abundance and mineralogy of sulfur on  \n<!-- image -->  \nradar-bright and low thermal emissivity highlands and in enigmatic geological formations such as steepside domes, the Venera 8 site, the canali, and tessera terrains are much less specific. The critical question to constrain sulfur abundances in non-basaltic crustal materials depends on whether Venus experienced oxidation of crustal and mantle materials, aqueous history, or plate tectonics ( Sect. 16.4 ).', '16.2.2.1  Mafic Igneous and Altered Materials': "Venera 13 and Venera 14 (1982) and Vega 2 landers (1985) provided information on rock-forming elements through the XRF analysis (Table 16.3). These measurements are the only data on sulfur content in surface and crustal materials. The composition of rock-forming elements in these probes and concentrations of U and Th in Vega 2 rocks (Table 16.4) suggest mafic silicate composition. Aside from bulk sulfur abundance, Venera 14 and Vega 2 rocks' major element composition is like that of olivine basalts that occur on oceanic islands, oceanic plateaus, and greenstone belts and are found in N-MORB tholeiite series (e.g., Surkov et al. 1984; Barsukov 1986d; Barsukov 1992; Kargel et al. 1993; Basilevsky et al. 2007; Treiman 2007; Filiberto 2014). A factor analysis of venusian and oceanic igneous rocks by Ivanov (2016) indicates a similarity of Venera 14 rocks with MORB and oceanic island basalts. Barsukov et al. (1986d) assigned the composition of Vega 2 rocks as olivine-gabbro-norite, the igneous rock common in Precambrian layered intrusions and Mesozoic ophiolite complexes. Potassium-rich Venera 13 composition suggests an alkaline mafic rock (e.g., leucitic basalt, olivine leucitite, phonolite) (Barsukov et al. 1982b; Barsukov 1992; Nikolaeva 1990; Kargel et al. 1993; Filiberto 2014; Treiman 2007). No terrestrial analogs for Venera 13 and Vega 2 rocks are inferred by Ivanov (2016). The large error bars for major elements do not allow for better analogs to be found among diverse groups of terrestrial mafic igneous rocks. Any definitive determination of the petrological type of these three samples is also limited by the need for data on Na and trace elements (Ni, Sr, Rb, REE, etc.).  \nLayered and porous rocks at the landing sites of Venera 13 (Fig. 16.1) and Venera 14 rocks could be presented by physically weathered lava flows, pyroclastic deposits, or airborne deposits of impactgenerated particles (Garvin et al. 1984; Basilevsky et al. 1985, 2004). The Magellan radar-based morphology of landing site regions (Weitz and Basilevsky 1993; Basilevsky et al. 2007) is consistent with basaltic composition. The bulk sulfur content in Venera 13 and Vega 2 samples is significantly higher than in typical terrestrial, lunar, martian basalts, and eucrites (~0.05 -0.3 wt%) and suggests exogenic ( Sect. 16.3.2 ), but not cosmic ( Sect. 16.4.3 ), sources of sulfur. Although high sulfur contents are  \n<!-- image -->  \nexpected in S(VI)-bearing magmas oxidized beyond the Ni-NiO (NNO) buffer (e.g., Baker and Moretti 2011), more reduced magmas with S(-II) are suggested for Venus' plain -forming basalts, as discussed below.  \nThe overall low dialectic constant (4 -6, Pettengill et al. 1988, 1992, 1997) of surface materials below ~6054 km radius and in landing site regions do not indicate abundant phases such as iron sulfides and/or oxides such as magnetite. The low dielectric constant (< 8) in older surface materials at lowlands, such as landing sites of Venera 13, 14, and Vega 2, may reflect unabundant iron sulfides/oxides in mafic bedrocks, products of their chemical alteration, and/or covering by silicate/sulfate-rich dust. A low sulfide content in S-rich Vega 2 materials is indirectly supported by in situ measured specific electrical resistance of ~10 6 Ωm , characteristic of terrestrial basalts at ~770 K. In contrast, the electrical resistance of Venera 13 and 14 materials (89 and 73 Ωm, respectively, Kermurdzhian et al. 1983) suggests h ighly conductive materials. However, these low electrical resistance values are inconsistent with an ordinary radar emissivity and overall low dielectric constant of these and other (Venera 8, 9, and 10, and Vega 1 and 2) landing regions reported by Garvin and Head (1986a, b) and Weitz and Basilevsky (1993). Abundant iron sulfides and/or oxides in Venera 13 and 14 materials are inconsistent with their iron content (Table 16.3), which is typical for mafic rocks. Therefore, the electrical resistance data for Venera  \n13 and 14 materials could be flawed.  \nThe regional geology of Venera 9, 10, 13, and 14 and Vega 1 and 2 landing ellipses (Weitz and Basilevsky 1993; Basilevsky et al. 2007) suggests a dominance of plains formed by high-yield eruptions of lowviscosity magmas (Head et al. 1992; Kargel et al. 1993; Crumpler et al. 2007; Ghail et al. 2024). The morphology, density, and hardness of rock fragments at these landing sites of Venera 9 and 10 are consistent with physically degraded lava fragments (Fig. 16.7) (Florensky et al. 1977, 1983; Garvin et al. 1984; Basilevsky et al. 1985), and the low albedo at visible and near-IR ranges (Golovin et al. 1983) supports mafic composition in both sites. In addition to the results of the XRF analyses (Table 16.3), the mafic composition of surface rock is suggested from concentrations of K, U, and Th at landing sites of Venera 9 and 10 and Vega 1 and 2 (Table 16.4). Specifically, K, U, and Th concentrations are like those in sub-alkaline basalts and gabbro (Taylor 1991; Basilevsky et al. 1992; Kargel et al. 1993; Treiman 2007). However, no perfect match exists, especially concerning U and/or Th. Nikolaeva (1995) demonstrated that all lowK Venus' rocks are significantly enriched in K, U, and Th compared to N-MORB.  Nikoilaeva  \n27  \n<!-- image -->  \n(1997) noted that Venus' rocks (Table 16.4, except Venera 8) have higher U content than island-ach tholeiitic basalts, and Venera 9 material is also enriched in Th.  \nThe morphology of geological features on vast plains seen in orbital radar images (Fig. 16.8) is consistent with basaltic volcanism (Barsukov et al. 1986a; Head et al. 1992; Crumpler et al. 1997; Ghail et al. 2024). The common occurrence of small (< ~500 m in height) cinder cones on vast plains suggests mainly effusive eruptions of volatile-poor mafic melts. Vivid lava morphology at volcanic centers (Fig. 16.9) is also consistent with basalts, though ~1 µm surface daytime emission (Mueller et al. 2008) suggests the ultramafic composition of a volcanic rise.  \nAlthough sulfur has not been considered in Venus igneous petrology models and discussions, terrestrial analogs and experimental solubility data could be invoked to assess the fate of igneous sulfur. As on Earth, the abundance and mineralogy of magmatic sul fur should depend on magma's redox state and Fe(II) content (Wallace and Carmichael 1992; Baker and Moretti 2011; Boulliung and Wood 2022, 2023; Simon and Wilke 2024). The cited works suggest that magmatic sulfur has a minimum solubility at 1.0 -1.5 log f O2 units above the quartz-fayalite-magnetite (QFM) buffer (near the Ni-NiO buffer). At the minimum solubility, an abundance of S(-II) equals that of S(VI). The solubility strongly increases at log f O2 < IW-3 and can reach a few wt% S (e.g., Kiseeva and Wood 2015; Namur et al. 2016) when sulfur could form melt complexes with Ca and Mg. Such a high solubility could characterize molten enstatite chondrites and aubrites (McCoy et al. 1999; Berthet et al. 2009) and mafic magmas on Mercury (Renggli et al. 2024, for a review) but not mafic melts that formed Venus' plains.  \nWänke et al. (1973) noted that the FeO/MnO ratio in terrestrial, lunar, and achondrite samples correlates with f O2 in the magmatic system. Schaefer and Fegley (2017) stated that the FeO/MnO ratio in surface probes (Table 16.3) suggests f O2 of corresponding melts between terrestrial upper mantle (~QFM buffer, Frost and McCammon 2008) and martian (~IW to IW+2 log f O2 units, Wadhwa 2008) values. However, FeO vs. MnO plots of Kargel et al. (1993) with error bars of XRF data do not indicate a difference between Venus' and terrestrial mafic rocks . The Earth-like (at and slightly below QFM) moderately reducing conditions do not imply different sulfur solubility in magma than in terrestrial counterparts. The redox state of venusian mafic igneous rocks implies Earth-like sulfide mineralogy (pyrrhotite, pentlandite, Fegley and Treiman 1992a) and ~0.1 wt% S (likely within 0.05 -0.2 wt%) in unaltered basalts (Moore and Fabri 1971; Wallace and Carmichael 1992; Kiseeva et al. 2024). Neither  \n<!-- image -->  \nmagmatic sulfates (e.g., anhydrite) nor Mg/Ca sulfides (oldhamite, niningerite) that could contribute to elevated sulfur content are expected in low-viscosity silicate melts forming plains and volcanic centers. Therefore, abundant sulfur in Venera 13 and Vega 2 samples could be in secondary sulfates, consistent with the expected chemical alteration pathways of Ca-Na-bearing basaltic glasses and Ca-rich pyroxenes at Venus' conditions ( Sect. 16.3.2 ).  \nMorphologically fresh crater outflows and corresponding air-born parabolic deposits of ejected materials (Basilevsky et al. 2004) are characterized by a slightly higher dialectic constant (~8, Campbell et al. 1992) that could be due to more abundant iron sulfides in geologically recently excavated materials. Magellan radar data of Pettengill et al. (1992) on ejecta from large craters (e.g., Mead, Stanton, Boleyn) also implies the delivery of materials with a high dielectric constant. This suggests higher abundances of corresponding minerals at depths of a km or more. The decimeter-wave emission data suggests more abundant iron sulfides and/or oxides below 0.7 -0.9 m beneath the layer with a dielectric constant of ~4.5 (Antony et al. 2022). However, the interpretation of radar data regarding materials with different dielectric properties is not unique. Physical structure (shape, size, or distribution of rock fragments) could affect microwave emissivity (Ford and Pettengill 1983; Pettengill et al. 1992), such as through multiple (volume) scattering (Tryka and Muhleman 1992; Pettengill et al. 1992, 1996; Campbell et al. 1999; Bondarenko and Kreslavsky 2018).", '16.2.2.2 Evolved Igneous Rocks and Venera 8 Site': "The one-modal hypsometric curve on Venus (Ford and Pettengill 1992) does not indicate the presence of low-density felsic continental crust in large regions (Grimm and Hess 1997). The lack of stratovolcanoes and arch structures that signify subduction of lithospheric plates on Earth suggests a deficiency of evolved plutonic (diorites, granodiorites, granites) and volcanic rocks (andesites, dacites, rhyolites). However, evolved igneous rocks formed via differentiation of mafic melts in magma chamber should be present. MELTS magmatic code models (Shellnutt 2013) show that phonolites and rhyolites could result from anhydrous and hydrous fractional crystallization of mafic magmas with the composition of Venera 13 and Venera 14 rocks. A significant differentiation is expected in large, long-standing magma chambers, as it occurred on Earth in Bushveld, Skaergaard, Stillwater, and other layered inclusions of bulk mafic or ultramafic composition within continental crust. Geophysical models for the venusian crust (Head and Wilson 1992) suggest a more common occurrence of shallow and large magma chambers  \n<!-- image -->  \nthan on Earth that reside in neutral bouncy zones. Slower solidification of such magma reservoirs in hot venusian crust favors igneous differentiation to felsic melts at the ch amber's tops and ultramafic cumulates at the base.", 'Steep-sided Domes': "Although most SiO2-enriched rocks could be preserved in crystallized plutons, some low-viscosity melts could have reached the surface from shallow magma batches. About 145 widely distributed pancakeshaped steep-sided domes (Fig. 16.10) could be formed by such melts (Fink et al. 1993; McKenzie et al. 1992; Moore et al. 1992; Parvi et al. 1992). These features are morphologically like rhyolitedacite domes but are 10 -100 times wider and have larger volumes (25 -3400 km3) orders of magnitude than terrestrial analogs (Parvi et al. 1992). The volume of steep-sided domes implies large magmatic chambers, as Head and Wilson (1992) suggested. The round morphology of the domes indicates that their formation in one eruption did not cause explosions, major degassing, and pyroclastic activity, which are suppressed at high ambient pressure (Head and Wilson 1986). The domes are commonly associated with coronae, which are unique circular tectonic-volcanic structures up to 1000 km in width supposedly formed above mantle plumes and large crustal magma reservoirs (e.g., Stofan et al. 1991, 1997; Head and Wilson 1992; O'Rourke and Smrekar 2018). If evolved melts are formed through closesystem differentiation in magma chambers, abundant sulfides are not expected in domes. On Earth, iron sulfides commonly remain dispersed in inner parts of closed system crystallizing mafic magma chambers, through sulfides of chalcophile elements (Ni, Cu) in association with pyrrhotite sometimes provide significant concentrations of Fe, Ni, and Cu in lower parts of magma chambers (Barnes et al. 2017; Latupov et al. 2024; Maier et al. 2018).", 'Are there any Massive Sulfide Deposits?': "The steep-sided domes are commonly associated with plains adjusted to tessera terrains ( Sect. 16.2.2.4 ). If tesserae are presented by evolved magmatic, metamorphic, or sedimentary materials, steep-side domes may reflect partial melting of SiO2-enriched country rocks around mafic magma chambers that reside in a zone on neutral buoyancy (Parvi et al. 1992). Although such processes are not rare within the continental crust (Eichelberger 1978; Glazner and Usslet 1988), a scarcity of domes within tesserae suggests a different pathway and massive sulfide deposits are not expected in domes.  \n<!-- image -->  \nMagmatic assimilation of crustal rocks could affect the fate of Venus ' sulfur in several ways. In one scenario, assimilation of putative crustal sedimentary sulfur (e.g., in anhydrite) would lead to a reduction of sulfur and the formation of massive sulfide deposits in magma chambers, as suggested for Norilsk-type intrusions (Gorbachev and Grinenko 1973; Grinenko 1985). Assimilation of sulfate-rich rocks increases magma's f O2 and sulfur solubility in S(VI) form. It could lead to magmatic sulfates (anhydrite) crystallization from oxidized silicic magmas, as it occurs in some subduction-zone magmatic systems (Masotta and Keppler 2015). In all cases, abundant sulfides are not expected in melts that form steep-sided domes. Suppose domes formed from bubble-rich melt accumulated in upper parts of basaltic magma chambers (Parvi et al. 1992); sulfur content and mineralogy of high viscosity features would not differ from those in basalts. Although sulfide vs. sulfate contents in evolved melts would constrain magma origin, the sulfur mineralogy of Venus' evolved magmatic compositions is unknown. The similarity of radar emissivity of domes with that of surroundings (Parvi et al. 1992) does inform the concentration of sulfides and/or Fe oxides because all surfaces at low and moderate elevations could be covered by silicate-dominated dust.  \nIt is unclear if the massive accumulation of magmatic sulfides in both magmatic chambers and thick flows even occurred on Venus. On Earth, the accumulation of massive sulfides in magma chambers is related to an environment of crustal sulfur that accumulated through non-magmatic processes (Grinenko 1985). Massive sulfide deposits are not common in layered mafic intrusions (Latupov et al. 2024; Maier et al. 2001; Barnes et al. 2017), and exceptions (e.g., Uitkomst intrusion) likely involve assimilation of country rocks (Maier et al. 2018). Sulfide deposits in lower parts of komatiite lava flows result from the assimilation of crustal sulfur originating from surface materials such as sediments (e.g., Bekker et al. 2009). If Venus bypassed an initial oceanic phase with an O2-bearing atmosphere (e.g., Zahnle and Kasting 2023, Sect 16.4.2 ), sedimentary sulfate reservoirs might not have formed nor participated in igneous activities. Conversely, Norilsk-like magmatic systems with extensive sulfide deposits could have emerged.", 'Venera 8 Site': "The concentrations of U, Th, and K measured in situ by Venera 8 (Table 16.4) lander suggest an evolved magmatic composition between mafic and ordinary silicic rocks. There are no images or spectroscopic data from the landing site, and the petrological type and geodynamic settings of Venera 8 rocks are subject to speculation.  There is no convergence among the proposed compositions of silicate rocks.  \n<!-- image -->  \nAndesitic, high-K calc-alkaline, leucitic, shoshonitic, melasyenitic, syenitic, nepheline syenitic, and quartz monzodiorite rock types have been proposed since 1973 to match concentrations of K, U, and Th (Nikolaeva 1990; Kargel et al. 1993; Shellnutt 2019, and references therein). Nikolaeva (1990) found a good match between terrestrial quartz monzodiorites and quartz syenites. Suggested compositions are like felsic to intermediate igneous rocks (diorite/granodiorite) typical of terrestrial convergent margins of lithospheric plates. All these terrestrial analogs were formed with significant involvement of crustal materials, which allowed Nikolaeva (1990) to suggest the formation of Venera 8 melts within an Earthlike continental crust. Geochemical considerations of magma fractionation by Nikolaeva and Ariskin (1999) ruled out the formation of Venera 8 melts through fractional crystallization of mafic melts.  They suggested that the Venera 8 composition could result from meeting an eclogite. In contrast, the petrological modeling of Shellnutt (2019) with MELS shows that Venera 8 rock compositions proposed by Nikolaeva (1990) could form via fractional crystallization of hydrous Venera 14 type mafic melts at f O2 of the Ni-NiO (QFM + 0.7) buffer. The silicate composition of Venera 8 materials inferred by Shellnutt (2019) corresponds to a magnesian, calc-alkalic trachydacite/granodiorite rocks common in Archean greenstone belts. A steep-sided dome of 23 km in diameter in a peripheral part of the Venera 8 landing region (Abdrakhimov 2001; Parvi et al. 1992) may consist of such rocks. Shellnutt (2019) suggested that Venera 8 landed on a bimodal volcanic complex formed from the compositionally evolved magma chamber. Magellan images and geological mapping indicate a high abundance of small volcanic shields in the landing ellipse (Abdrakhimov 2001; Basilevsky et al. 1992; Weitz and Basilevsky 1993; Basilevsky 1997; Basilevsky et al. 2007). The size (a few km across) and gentle morphology of volcanic shields suggest low rates of magma emplacement that could have occurred from widespread, isolated, longlived, and possibly shallow magma reservoirs (Head et al. 1992; Crumpler et al. 1997). These fields of small shallow shields are inconsistent with silicic viscous melts but could have been formed by mafic lavas analogous to alkaline oceanic island hotspots. Shield plains occupy about 10 -15% of Venus' surface and are partially buried by plains with wrinkle ridges, a dominant surface type on lowlands (Ivanov and Head 2004). Therefore, Venera 8-type igneous rocks could represent a significant portion of the upper crust. Although the abundance and mineralogy of sulfur have not been discussed in the Venera 8 rock, its probable mafic alkaline composition cautiously suggests a sulfur fate like its terrestrial counterparts.  \n<!-- image -->", '16.2.2.3 Canali-type Channels': "Unusual lava channels (canali) and more complex channels and valleys are observed primarily on volcanic plains (Baker et al. 1992, 1997; Komatsu et al. 1992, 1993) (Fig. 16.11), though rare canali are seen in tessera terrains (Fig 16.12). Some features are outflow channels related to impact craters. Rare valley networks suggest supping from a shallow subsurface. Although most of the ~200 channels are tens to hundreds of km long, one canale, Baltis Vallis, is almost 7000 km long. The channels are typically characterized by widths (1 -3 km) and large apparent width-to-depth ratios that remain nearly constant along the length. The bottom of the Baltis Vallis is 20 -100 m lower than the surrounding plains (Oshigami and Namiki 2007). Canali channels have meanders, but cut-off meanders, tributaries, distributaries, and levee structures are rare. Their morphology suggests erosion by low-viscosity melts supplied from subsurface reservoirs through a single main conduit, maintaining steady effusion rates. Both thermal and mechanical erosion by highly turbulent lava was discussed by Baker et al. (1997), Williams-Jones et al. (1998), and Oshigami and Namiki (2007). The morphology of canali does not indicate significant crystallization of melts at ambient conditions at the time of impeachment, which constrains melt composition and past climate. In this context, silicate lava compositions are less favorable than non-silicate low-temperature melts, which do not solidify readily during emplacement.  \nMicrowave or other remote sensing data provide little compositional information on melt composition in canali and valleys. Suggested low-viscosity silicate melts included ultramafic (komatiites) and alkaline mafic (picrites, olivine nephelinites) compositions (Baker et al. 1992; Komatsu et al. 1992; Gregg and Greeley 1993; Williams-Jones et al. 1998). The primary concern is their rapid cooling, which does not allow the formation of prolonged morphologically uniform channels (Komatsu et al. 1992). The thermal erosion by hot ultramafic melts could be inconsistent with their rapid cooling (Gregg and Greeley 1993, 1994; Williams-Jones et al. 1998) and topography profiles across Baltis Vallis (Oshigami and Namiki 2007).  On Earth, ultramafic melts commonly assimilate sulfur from crustal and surface rocks, as evidenced by sulfur isotopes and other data (Arndt et al. 2008; Bekker et al. 2009; Kubota et al. 2022). Bottoms of some komatiite lava flows and sills reveal massive sulfide deposits possibly formed by incorporating sulfides from underlying rocks affected by hot lava (Lesher 1989; Arndt et al. 2008; Bekker et al. 2009). Likewise, ultramafic lava on Venus could have affected sulfides and/or sulfates accumulated on volcanic plains through gas-solid reactions ( Sects.  16.3.2, 16.4.4 ) and/or fumarolic activity. Whether ultramafic melts assimilated crustal sedimentary sulfur depends on aqueous ( Sect. 16.4.2.3 ) vs.  \n<!-- image -->  \nanhydrous geological history. Regardless of the sulfur source in ultramafic melts, putative bottom sulfide deposits may not be exposed due to the minute erosion and surface modification rates since the lava emplacement (Arvidson et al. 1992; Kreslavsky and Bondarenko 2017; Carter et al. 2023). Indirect detection of subsurface sulfide deposits in canali and valleys revealed by penetrating radar on the EnVision Venus orbiter (Ghail et al. 2018, Sect. 16.5.2 ) would be consistent with ultramafic melts.  \nDiscussed non-silicate melts are sulfur (Baker et al. 1992; Williams-Jones et al. 1998), carbonatite (Baker et al. 1992; Kargel et al. 1994); carbonatite-chloride (Williams-Jones et al. 1998), carbonate-sulfate (Kargel et al. 1994; Treiman 1995, 2009), and chloride (Zolotov and Mironenko 2009; Zolotov 2019) melts. The eutectic temperature of those salt mixtures is significantly lower than that of silicates, and some complex salt melts could be stable under current ambient conditions. These parameters are more consistent with mechanical erosion rather than thermal erosion by hot silicate melts (Oshigami and Namiki 2007).  \nNative sulfur melts at 388 K at 1 bar. Sulfur lava flows are rare on Earth (Harris et al. 2000) and are suggested to occur on Io (Greeley et al. 1984; Kargel et al. 1999). Formation of native sulfur through incongruent melting of pyrite and/or pyrrhotite was suggested for Io by Lewis (1982) and is considered possible on Venus (Baker et al. 1992). It is unclear if Venus ' crustal conditions allow the production of abundant sulfur from sulfides, which is not common on Earth. Although liquid sulfur becomes more viscous upon cooling from a magmatic (~1000 K) to Venus' surface temperature, it is still less viscous than ultramafic (komatiitic) melt.  At current surface conditions, liquid sulfur does not boil but is subjected to rapid evaporation after emplacement due to high vapor pressure (~1.5 bar at 750 K), especially if it erupts at a temperature above ~750 K. Evaporation of liquid sulfur in canali suggests decreasing discharge and widths of flows with increasing distance from a source, which is not observed.  \nCarbonatites are regarded as suitable analogs for canali's lavas due to their lower viscosity compared to silicate melts and because they are associated with alkali mafic and ultramafic rocks, which may form through a low-degree partial melting of a carbonated mantle peridotite (Baker et al. 1992; Kargel et al. 1994; Treiman 1994). Mechanical erosion by low-viscosity carbonatite melts is consistent with the morphology of channels and cooling models of canali's melts (Williams -Jones et al. 1998; Oshigami and Namiki 2007). Although typical carbonatites consist of Mg-Ca-Na-K carbonates, sulfates are usually minor components with chlorides, phosphates, and silicates. Alkali-rich carbonatites have lower  \n34  \n<!-- image -->  \nviscosities and lower solidus temperatures than Ca-Mg carbonatites (Williams-Jones et al. 1998). Ternary Ca-Na-K carbonate system with admixtures have eutectic between 850 and 938 K, and natrocarbonatite lavas of Oldoinyo Lengai volcano (Tanzania) have eruption temperature of ~770 -780 K (Dawson et al. 1995) only slightly hotter than Venus ' ambient temperature. These low melt temperatures could reflect an admixture of alkali chlorides, and alkali-carbonate-chloride melts were considered suitable valleyformed melts (Williams-Jones et al. 1998). Kargel et al. (1994) presented arguments for carbonatitesulfate melts that could form by melting crustal carbonate and sulfate deposits or a carbonated mantle.  \nA common occurrence of carbonatites and carbonate-sulfate melts is questionable because of several factors. The issues include the high eutectic temperature (1250 K, Treiman 1995) and high viscosity of the CaCO3-CaSO4 system (Kargel et al. 1994), unlikely formation of crustal carbonate and sulfate deposits without surface liquid water ( Sect. 16.4.2 ), doubtful highly oxidized upper mantle with magmatic S(VI) ( Sects. 16.4.1, 16.4.2 ), and thermal decomposition of crustal carbonates upon heating caused by runaway greenhouse or greenhouse caused by global volcanic degassing ( Sects. 16.4.2; 16.4.4.1 ). Sulfate-carbonate melts have not been reported on Earth. The atmospheric CO2 mass exceeds that of terrestrial crustal carbonates ( Sect. 16.4.3 ), and abundant carbonated materials are not expected in the interior of Venus.  Finally, the instability of Ca, Na, and K carbonates with respect to atmospheric SO2, HCl, and HF ( Sect. 16.3.2.2; Zolotov 2018) suggests chemical weathering of carbonatites. No carbonates could be detected with remote sensing methods that probe the surface materials.  \nChemical models for a reduced early Venus water ocean suggested the subsequent formation of CaCl2NaCl evaporites and their involvement in shallow subsurface melting (Zolotov and Mironenko 2009; Zolotov 2019; Sect. 16.4.2 ). Low-viscosity Ca-Na chloride melts with a eutectic temperature of 786 K (Tian et al. 2021) could be responsible for the formation of canali, crater outflow valleys, valleys with chaotic terrains, and valley networks related to sapping (Zolotov 2019). The greenhouse warming caused by volcanic degassing during global volcanic resurfacing (e.g., Solomon et al. 1999) could have caused the melting and mobilization of crustal chloride deposits. The superior capacity of Ca-Na-Cl melts to dissolve silicates could have contributed to the erosion and the development of the unique morphological characteristics of canali. The presence of sulfates in chloride-rich bedrocks in canali and valleys would indicate aqueous early history and oxidized surface water ( Sect. 16.4.2 ). However, the instability of Ca and Na chlorides with respect to atmospheric SO2 ( Sect. 16.3.2.2 ) suggests the formation of sulfate weathering crust on chloride rocks if that are present in canali and outflow valleys.  \n<!-- image -->", '16.2.2.4 Tessera Terrains and Highland Rocks': "Tessera terrains (Figs. 16.12, 16.13) are presented by pervasively tectonically deformed complexes that occupy about 8% of Venus ' surface area (Barsukov et al. 1986a; Bindschadler and Head 1991; Ivanov and Head 1996, 2011). They are elevated massifs of ~10 2 -10 3 m km in size, and some tessera terrains occupy fractions of highland plateaus. More miniature, isolated tessera terrains are surrounded and embayed by materials of volcanic plains. Although impact crater density on tesserae is statistically indistinguishable from that of surrounded ridged and densely lineated plains (Kreslavsky et al. 2015), stratigraphic relationships suggest tesserae as the oldest terrains on the visible part of the geological history (Ivanov and Head 2011, 2015). However, an apparent improvement of some adjoining geological units in tessera-forming tectonic deformations suggests overlapping formation times of tesserae and other heavily tectonized terrains (densely lineated and ridged plains, ridge belts) (Ivanov and Head 2015), consistent with overlapping retention ages of impact craters (Kreslavsky et al. 2015). Photogeological analysis of Magellan radar images by Ivanov (2001) showed that tessera precursor material appears as plains and that some tesserae could have the same (basaltic) composition as adjusted plains. A mafic composition of tessera composed of deformed thick lava formations was  \nhypothesized in some tectonic interpretations (e.g. Hansen 2006).  \nThe eroded layered and folded rock structures on Magellan radar images of several tesserae (Fig. 16.12, 16.13) suggest their sedimentary and/or volcanic origin.  Linear structures on radar images could reflect deposition through actions of impacts, wind, or liquid water. Extended parallel linear features imply the mafic/ultramafic composition of lava flows and pyroclastic deposits in a case of volcanic origin. Large intrusive complexes (granitoids, anorthosites, layered mafic massifs) are inconsistent with these structures (Byrne et al. 2021).  \nThe geomorphology of plains within tesserae is like that of regional plains and suggests basaltic composition (Ivanov 2001; Gilmore and Head 2018), though some radar-dark areas in local lows could be products of aeolian deposition of fine-grained material (Byrne et al. 2021). The basaltic lava composition of intratessara plains is consistent with only one steep-sided dome (Parvi et al. 1992) and no signs of pyroclastic deposits, exploded calderas, or stratovolcanoes. In contrast to preceding morphology-based works, the lava complex within the tessera highlands of Ovda Regio (Fig. 16.12) is  \n<!-- image -->  \nsuggested to be basaltic in composition based on the interpretation of Magellan radar images regarding surface roughness (Wroblewski et al. 2019).  \nSome interpretations of tessera morphology, deformation patterns, high gravity-topography ratios, and geodynamic models did not exclude the silicic composition of highland plateaus in supposedly exposed fragments of putative felsic continental crust (Romeo and Capote 2011; Resor et al. 2021). The interpretation of gravity-topography data by Maia and Wieczorek (2022) suggests an Airy isostatic regime for highlands and a 15 -40 km thick crust is not indicative of crustal composition. However, the viscous relaxation models of Nimmo and Mackwell (2023) show that quartz-dominated compositions relax too quickly to be plateau-forming materials. A lower dielectric permittivity on a tessera than on plains suggests a lower density of surface materials due to felsic and/or physically weathered rocks (Kreslavsky et al. 2000).  \nInterpretations of nighttime ~1 µm thermal emissivity of surface materials from Galileo (Hashimoto et al. 2008) and Venus Express (Mueller et al. 2008; Helbert et al. 2008; Basilevsky et al. 2012; Gilmore et al. 2015) do not exclude more felsic tessera compositions than materials on plains with basaltic lava morphology. Thermal surface emission in the visible range obtained during the Parker Solar Probe flyby agrees with near-IR data and tentative compositional interpretations (Wood et al. 2021; Lustig-Yaeger et al. 2023). Experimental data on the thermal emission of terrestrial materials at Venus' temperatures (Dyar et al. 2020, 2021; Helbert et al. 2021; Treiman et al. 2021) suggest relatively Fe(II)-poor surface silicate materials on tessera highlands. However, the explanations of the thermal emission from highlands in terms of composition need confirmation because, in addition to lower surface temperature, the apparent low thermal emission from high-altitude surfaces could be due to physical properties (e.g., grain size) and lower actual temperatures than inferred from Magellan altimetry (Basilevsky et al. 2012; Gilmore et al. 2015, 2017). Another possible reason for low emission is coating by light-toned products of chemical weathering that could have formed during the formation of volcanic plains (Gilmore et al. 2015; Sect. 16.4.4 ) or in the current epoch as more stable phases (e.g., Ca and Na sulfates) on highlands (Zolotov 2018, Sect. 16.3.2.2 ). If low thermal emission from tessera highlands reflects the composition of light-toned bedrock rather than a coating, layered features (Fig. 16.13; Byrne et al. 2021) could be ancient sediments that are depleted in ferrous silicates and iron sulfides. Putative sedimentary rocks could be sandstones, clays, or shells, and chemical sediments could be carbonates, phosphates, and saltrich evaporitic deposits (chlorides, sulfates). Although felsic silicate deposits could be aeolian (loess) or  \n37  \n<!-- image -->  \nimpact in origin, the formation of abundant felsic rocks likely required aqueous processes that favored granitization (Campbell and Taylor 1983; Taylor and McLennan 2008). Formation of the listed chemical sediments required liquid water that could have occurred on early Venus ( Sect. 16.4.2 ). Whether sulfur is present in sulfates depends on the fate of oxygen that was strongly affected by hydrogen escape ( Sects. 16.4.1 , 16.4.2 ). Detection of sulfate-rich layered rocks in tessera with following missions ( Sect. 16.5 ) would indicate past aqueous and oxidizing conditions caused by hydrogen escape from a moist greenhouse atmosphere ( Sect. 16.4.2 ).", '16.2.2.5  Low Radar Emissivity Highlands': "At Venus' highlands, roughly above a 6054 km radius, low microwave emissivity (high reflectivity) suggests a presence of electrical semiconductor material with a high dielectric constant (Pettengill et al. 1982, 1988, 1992; Simpson et al. 2009) that might presented by sulfides. The low-emission intensity pattern varies with elevation and location, though local geology and geomorphology have no clear effect (Arvidson et al. 1994; Klose et al. 1992; Treiman et al. 2016), Fig. 16.14. The data suggest the dielectric constant of 50 or more in surface materials (Pettengill et al. 1988); through neither composition, no origin of low-emissivity materials is known. Other interpretations of these data include ferroelectric minerals (Shepard et al. 1994) and volume scattering unrelated to the composition (Tryka and Muhleman 1992; Pettengill et al. 1992, 1996, 1997; Bondarenko and Kreslavsky 2018).  \nBoth physical and chemical processes have been invoked to explain the low-emissivity phenomenon on highlands. Dense high-dielectric phases (magnetite, iron sulfides) could have been concentrated in lag deposits by blowing low-density grains by wind (Greeley et al. 1991). Although iron sulfides are commonly invoked to explain the low radar emissivity at highlands (Pettengil et al. 1982; 1988; Klose et al. 1992; Wood 1997), the thermodynamic stability of iron and other sulfides may not be correctly constrained because of gas-gas chemical disequilibria at elevations (e.g., Zolotov 2018). Nevertheless, pyrite could be more stable than magnetite at elevations (Klose et al. 1992; Zolotov and Volkov 1992; Wood 1997; Zolotov 2018; Sect. 16.3.2.3 ), and a metastable existence of primary and/or secondary pyrrhotite is not entirely excluded (Zolotov 2018).  Experimental works on the stability of metal sulfides at simulated conditions of Venus surface ( Sect. 16.3.2.3 ) need to provide a conclusive understanding of their stability and contribution to the low radar emissivity.  \n<!-- image -->  \nConsiderations of gas-solid type chemical equilibria do not exclude high-altitude condensation of heavy metal sulfides such as PbS and Bi2S3 (Schaefer and Fegley 2004) at highlands ( Sect. 16.3.2.3 ). However, the spatially variable low-altitude boundary of low-emissivity highlands is inconsistent with condensation from the atmosphere, which could cause a stricter transition altitude through the globe. The increase in radar emissivity at the highest elevations (Klose et al. 1992; Arvidson et al. 1994; Treiman et al. 2016; Fig. 16.14) disagrees with condensation above certain altitude-temperature levels.  \nIn addition to compounds with an elevated dielectric constant, materials with ferroelectric properties have been proposed to explain radar brightness at highlands and decrease in the brightness at some of the highest elevations (Arvidson et al. 1994; Shepard et al. 1994; Brackett et al. 1995). Radar emissivity of the Ovda Regio highland region (Fig. 16.12) does not contradict the occurrence of a ferroelectric phase at higher elevations that becomes paraelectric at lower altitudes at higher temperatures. Although several such ferroelectric phases have been proposed (e.g., chlorapatite, Treiman et al. 2016), the problem still needs to be solved, and no S-bearing minerals are among them.", '16.2.3 Bulk Venus, its Core, and Mantle': "Although Venus ' bulk composition and bulk S content are unknown, inferences could be obtained from Venus' size and mean density, the Earth's composition, and the composition of Venus' surface and the atmosphere. Most estimations are model-dependent and reflect physical-chemical processes on early Venus ( Sect. 16.4.1) . Physical models for planetary accretion suggest the formation of Earth and Venus from compositionally similar materials ( Sect. 16.4.1 ). The similarity of bulk planetary compositions is supported by comparable densities on the planets, by roughly Earth-like abundances of K, U, and Th in five Venera and Vega probes of surface materials (Table 16.4; Sects. 16.2.2.1, 16.2.2.2 ), by the widespread volcanic features with basaltic lava morphology ( Sect. 16.2.2.1 ), by a matching of rockforming elemental composition of sampled surface rocks (Table 16.3) with terrestrial igneous counterparts ( Sect. 16.2.2.1 ), and by the similarity of masses of carbon and nitrogen in upper envelopes of the planets ( Sect. 16.4.3 ).  \nStrictly identical compositions of Venus and Earth would result in ~1.9% greater Venus' density than observed (BVTP 1981). However, the pressure-corrected density of Venus is ~3% less than that of Earth (Phillips and Malin 1983). The difference is commonly attributed to lower iron and sulfur content and  \n<!-- image -->  \nincreased oxygen abundance on Venus. Higher volatility of sulfides in hotter parts of the solar nebula could have contributed to slightly lower sulfur content in Venus-forming bodies (BVTP 1981). However, this mechanism is inconsistent with S-rich materials on the surface of Mercury (Nittler et al. 2018; Renggli et al. 2024) and with the likely formation of terrestrial planets from planetary embryos and differentiated planetesimals with Fe-S cores ( Sect. 16.4.1 ). The stochastic accretion of compositionally diverse planetary embryos, including different core sizes and variable degrees of impact stripping, could have resulted in various planetary densities and lower bulk Fe and S contents on Venus. A higher basalt/eclogite ratio due to the warmer upper interior could also contribute to the low density of Venus (BVTP 1981). Based on current data, it is impossible to make a definitive prediction regarding the bulk sulfur content, which primarily reflects the composition and size of the core.  \nIn early models for the bulk composition of terrestrial planets, the bulk sulfur content was determined by the abundance of sulfur in the core and the core's mass fraction. The latter was constrained by the bulk density of a planet and by the Birch-Murnaghan equation-of-state for planetary materials to approximate their densities at high pressures (Table 16.5). Mantles and crusts were considered S-free. The Ve2 model (BVTP 1981) was based on the results of equilibrium condensation calculations of solar nebular materials and is modified using feeding zones. That model has a sizeable S-rich core and a lowdensity mantle with elevated pyroxene/olivine and Mg/Fe ratios. Anders (1980) used either the solar S/K ratio or the amount of iron available after forming Fe-Ni metal and silicates in the solar nebula to constrain sulfur abundance in Venus (Ve3 model). The latter model remains commonly used for over 40 years (e.g., Xiao et al. 2021). Models Ve1, Ve2, and Ve3 have no oxygen in the core. In contrast, the pyrolite Ve4 model is characterized by a small S-poor and O-rich core with an olivine-rich and FeO-rich mantle. The Ve5 model represents a Fe-deficient composition for the low Venus density. The model has a low-density S-poor and O-rich core. Compared to similar models for the Earth (BVTP 1981), these assessments reveal slightly smaller mass fractions of the core and lesser sulfur abundances in the core and the whole planet. In agreement with these models, Aita et al. (2012) demonstrate that 4.8 wt% S in the core of Venus is insufficient to explain core density, and other light elements are needed. Recent models suggest S and Si as major light elements in cores on Earth, Mercury, Mars, and the Moon (e.g., Huang et al. 2019, Terasaki et al. 2019; Steinbrügge 2021), and Venus (Xiao et al. 2021; O'Neil 2021). However, the challenge is to discriminate S and Si in core compositions on all terrestrial planets.  \n40  \n<!-- image -->  \nThe range of core sulfur content of 1 -10 wt% corresponds to 0.3 -3 wt% of sulfur in the bulk planet, respectively (Table 16.5). The latter values are less than sulfur abundances relative to rock-forming elements (e.g., Si, Fe, Mg) in the solar photosphere, in Ivuna-type carbonaceous chondrites (CI) that best represent solar system abundances in the meteorite collection (Lodders 2021), and in returned samples of Ryugu carbonaceous asteroid (Yokoyama et al. 2022). Trapping of H2S gas to sulfides of Fe, Ni, and Mn in the solar nebula (e.g., Lauretta et al. 1997) could have been incomplete in the inner solar system. The deficiency of sulfides may also reflect their thermal vaporization in the inner solar system (Morgan and Anders 1980). These pathways agree with below solar S/(Si, Fe, Mg) ratios in ordinary and enstatite chondrites (Lodders and Fegley 1998) that could be significant building blocks of planetary embryos that formed Venus. Further indepth discussion on S in Venus' core is Sect. 16.4.1 .  \nThere are no identified or suspected samples of Venus rocks in the meteorite collection, and the elevated S content in otherwise basaltic materials in surface samples (Table 16.3) could be due to chemical alteration at the surface ( Sect. 16.3.2 ). A similarity of Fe content of these probes with terrestrial basalts and only a moderately more reduced Venus' mantle compared to the terrestrial counterpart suggested from the MnO/FeO ratio in surface samples ( Sect. 16.2.2.1 ) implies Earth-like sulfide-saturated mafic melts with 0.05 -0.2 wt% S and their mantle source regions containing 100 -300 ppmw (µg/g) S (Kiseeva et al. 2024, for a review).  \nIn the mantle transition zone on Earth, some (e.g., Ni-poor) sulfides could be liquid (Kiseeva et al. 2024). The hotter Venus ' geotherm in the crust and upper mantle would favor sulfide melting in the upper mantle or lower crust, and the zone of possible sulfide melts could be closer to the surface and possibly more comprehensive. Significant sulfide acclimated and trapped at the late stage of magma ocean crystallization ( Sect. 16.4.1 ) may affect the physical properties (viscosity, density, etc.) of the upper interior and concentration of chalcophile elements in magmas.", '16.3.1 Atmospheric Processes': 'The current understanding of atmospheric chemical and physical-chemical processes is based on photochemical and thermochemical modeling of selected reactions, chain reactions, and large multicomponent atmosphere segments. Major chemical processes involve S-bearing gases and  \n41  \n<!-- image -->  \ncondensates. Sulfur plays a key role in cloud chemistry and the mesosphere. The sulfur cycle has deserved persistent attention since the discoveries of H2SO4 in the cloud aerosol and SO2 gas above and beneath clouds in the 1970s. In preparation for the Pioneer Venus mission, Prinn (1975, 1978) made influential assessments of sulfur chemistry throughout the atmosphere. In the 1980s, the models were aimed at the interpretation of compositional data obtained with the Pioneer Venus, Venera 11 -15, and Vega spacecraft ( Sect. 16.2.1 ) in terms of photochemical and thermochemical processes and atmospheric dynamics (Winick and Stewart 1980; Krasnopolsky and Parshev 1981a, 1981b, 1981c, 1983; Yung and DeMore 1982; DeMore et al. 1985). Since the early 1990s, modeling efforts have targeted understanding the fates of S-bearing gases observed through ground-based, orbital, and spacecraft flyby data.  Significant understanding has been obtained with comprehensive models that consider the kinetics of photochemical and thermochemical reactions in the C-O-S-Cl-N system, mass balances, and eddy transport in the middle and upper atmosphere (Mills 1998; Mills and Allen 2007; Zhang et al. 2010, 2012; Jessup et al. 2015; Krasnopolsky 2012). Krasnopolsky and Pollack (1994) assessed the physical chemistry of the S-O-H system in the clouds in conjunction with the lower atmosphere. Yung et al. (2009) and Krasnopolsky (2007, 2013) specifically quantified sulfur in the lower atmosphere. Bierson and Zhang (2020) first modeled the atmosphere from 112 km to the surface, and further such modeling was performed by Dai et al. (2024). Stolzenbach et al. (2023) performed the first three-dimensional simulations of the coupled H-C-O-S-Cl system at 40 -95 km. Atmospheric sulfur cycles have been discussed by von Zahn et al. (1983), Prinn (1985a, 1985b), Esposito et al. (1997), Fegley et al. (1995, 1997a), Mills et al. (2007), Krasnopolsky (2007, 2013), and Bierson and Zhang (2020). Sect. 16.3.1.1 overviews chemical and physical processes that govern atmospheric sulfur. Sects. 16.3.1.2 to 16.3.10 provide details on the formation, transport, and consumption of S-bearing compounds, and atmosphere-surface interactions are described in Sect. 16.3.2 .', '16.3.1.1 Major Pathways and Sulfur Cycle': "At and above the cloud top, the fate of chemical species is governed by photochemical processes.  \nThermochemical reactions dominate in the middle and lower clouds and from the cloud deck to the surface. Transport and mixing occur in the lower and middle atmosphere through eddy diffusion and a global Hadley-cell type circulation that includes low-latitude upwelling and high-latitude downwelling, and polar and typography upwellings are possible. Both eddy diffusion and wind transport occur in the  \n42  \n<!-- image -->  \nupper atmosphere. Photochemical dissociation of O-bearing gases (CO2, SO2, H2O, OCS) at and above cloud tops creates oxidizing conditions that allow oxidation of SO2, OCS, and H2S to S(VI) species such as SO3 and H2SO4.  In the largely thermochemically controlled lower atmosphere below ~30 km, conditions are more reducing, S(VI)-bearing gases are absent, SO2 could be less abundant than at mid-30s km, and OCS is much more abundant than at higher altitudes (Table 16.1). Within and around clouds, the behavior of chemical compounds is affected by competition of photochemical and thermochemical reactions, mixing of species of photochemical and thermochemical origin via eddy diffusion and plumes, condensation-evaporation of sulfuric acid and possible polysulfur aerosols, dissolution-precipitation of mineral species such as metal sulfates, and by alteration of inorganic (Fe-metal, FeS, Mg silicates, etc.) and organic matter of cosmic origin. Short-term (e.g., hourly, daily) compositional variability of SO2 and SO at cloud top and above reflects solar insolation, rates of reactions, and competitive mixing. Causes of the long-term compositional variability of SO2 at cloud tops (Fig. 16.3) could be related to changes in atmospheric circulation patterns below and within clouds. The Hadley-cell circulation could account for latitudinal gradients of SO2, OCS, and CO observed at 30 -40 km. Observed and suspected vertical compositional gradients reflect changes in temperature, pressure (Seiff et al. 1985), and competitive rates of chemical reactions and gas mixing. In the deep lower atmosphere, rapid gas-phase thermochemical reactions are mainly compensated by reverse reactions, and differences in reaction rates account for the net production or loss of gases. In contrast to earlier chemical equilibrium models for the middle and lower atmosphere (Florensky et al. 1978; Oyama et al. 1980), the apparent lack of chemical equilibration between many gases above, within, and below clouds implies ongoing chemical reactions throughout the atmosphere. Thermochemical equilibration between major chemically active gases (CO2, SO2, OCS, and CO) is suspected only at the surface (e.g., Krasnopolsky and Parshev 1979; Fegley et al. 1997b; Krasnopolsky 2013, Table 16.2) but data on the gas composition (Table 16.1) are insufficient to confirm it. Indeed, there is no thermochemical equilibration between SO2, OCS, CO, and CO2 above a thin layer close to the modal Venus' radius (Krasnopolsky and Pollack 1994; Zolotov 1996; Fegley et al. 1997b), though some compounds (e.g., Sn species, H2SO4·nH2O(l) and H2SO4(g), Fig 16.5) could equilibrate at certain altitudes and environments.  \nSurface rocks and minerals ( Sect. 16.3.2 ), volcanic degassing, and space materials ( Sect. 16.4.3 ) provide net sulfur supplies to the atmosphere. The net supply of reduced gases (OCS, H2S, and S2) through oxidation of pyrite (FeS2) discussed by von Zahn et al. (1983), Prinn (1985a), and Fegley et al. (1995)  \n43  \n<!-- image -->  \nremains vague because of the uncertain composition of the deep atmosphere and possible stable existence of pyrite (FeS2). However, oxidation of pyrrhotite (Fe1-xS) by CO2 and H2O is likely. Although volcanic delivery of SO2, OCS, S2, H2S, and S-Cl gases could account for sulfur atmospheric inventory, current fluxes are ambiguous. A difference between supposed volcanic and atmospheric gas compositions implies the chemical consumption of volcanic gases after degassing (Wilson et al. 2024). For example, abundant OCS, S2, and Cl-S compounds in volcanic gases could be oxidized to SO2 ( Sect. 16.4.3 ). S-bearing gases could be released through the decomposition of sulfides in crustal and surface materials ( Sect. 16.3.2.3 ). Although the effect of space sources on sulfur inventory in the atmosphere is negligible ( Sect. 16.4.3 ), space materials could affect the composition of clouds (Zolotov et al. 2023; Mogul et al. 2024) and the mesosphere, as discussed below. In addition to net sources, S-bearing gases could be released into the atmosphere as products of gas-solid reactions at the surface ( Sect. 16.3.2 ).  \nThe net sink of atmospheric sulfur occurs through the formation of Ca and Na sulfates in exposed surface materials and a potential formation of secondary sulfides ( Sect. 16.3.2 ). Volcanic degassing and atmosphere-surface interactions determine the bulk mass of atmospheric sulfur. However, much lower rates of gas-solid reactions than gas-phase reactions exclude the effects of atmosphere-surface reactions on the short-term behavior of atmospheric gases and their altitudinal and meridional gradients.  \nThe atmospheric sulfur cycle is coupled with C, O, H, Cl, N, and Fe cycles that involve reactions between SO2, SO, SO3, Sn, OCS, CO, CO2, CS, CS2, O, O2, H2O, H2SO4, H2S, HS, ClO, SxClyOz, SxCly, NO2, and other gases. The coupling of S, C, and O cycles occurs through the oxidation of S-bearing species in upper clouds and interactions between OCS, SO3, CO, and Sn in the lower atmosphere. Sulfur and chlorine cycles are linked because, as in Earth's stratosphere, Cl-bearing species play roles as catalysts in upper clouds and the mesosphere, serve as intermediate species in reaction chains, and participate in forming and decomposing S-Cl-bearing gases. HCl gas delivered to upper clouds from the lower atmosphere is photochemically oxidized to Cl-O compounds that slow the net recombination of O and CO to CO2 and possibly participate in a series of reactions that produce Sn species, SxClyOz, and SxCly gases. The net effect of photochemistry in the middle atmosphere is the consumption of CO2, SO2, H2O, and HCl and the return of CO, H2SO4, and SO2Cl2 to the sub-cloud atmosphere (Krasnopolsky 2013). The net effect of thermochemical reactions in the lower atmosphere is the consumption of CO, H2SO4, and trace Clbearing gases to OCS, SO2, Sn, and HCl.  \n44  \n<!-- image -->  \nAlthough the critical processes in the atmospheric sulfur cycle (Fig. 16.15) are established, the roles of chemical and physical processes and reaction pathways are debated and remain uncertain. The following description of major pathways mainly reflects the results of atmospheric modeling by Yung and DeMore (1982), Mills (1998), Mills and Allen (2007), Zhang et al. (2012), Krasnopolsky (2007, 2012, 2013), Bierson and Zhang (2020), Pinto et al. (2021), and Dai et al. (2024). The details and other references are provided in the following sections.  \nReduced S-bearing gases (OCS, Sn, H2S, HS, CS, and CS2) form in thermochemical processes in the deep sub-cloud atmosphere where SO2 is the more abundant. SO2 and reduced gases are transported upward toward clouds through eddy diffusion and upwelling. The Hadley cell circulation (Fig. 16.16) leads to the low-latitude upwelling, that is supported by meridional gradients of SO2 (Marqc et al. 2021; Oschlisniok et al. 2021), OCS (Marcq et al. 2008, 2023), and CO ( Sect. 16.2.1 ). Local topography-related (Marcq et al. 2020, 2023; Lefèvre et al. 2020) and polar upwellings (Oschlisniok et al. 2021; Marcq et al. 2023) are suspected. OCS is consumed through interactions with atomic S gas at upper 20s km and SO3 at mid-30s km. Sn gases from below could condense in the lower clouds to form sulfur aerosols, Sx. A debated formation of Sn species through a chain of reactions in upper clouds may also contribute to Sx aerosols. Photochemical dissociation of CO2  \nCO2 + hv → CO + O (16.1)  \nO + O + M → O 2 + M (16.2)  \nin the mesosphere, a slower CO → CO2 back conversion creates an excess of O and O2 involved in oxidizing S-, C-, Cl-, N-, and Fe-bearing species at the cloud top and above. SO2, OCS, and H2S are consumed through photolysis and reactions with O and O2. The net process characterizes the formation of SO3 at the cloud top,  \nSO2 + CO2 → SO 3 + CO (16.3)  \nin which SO2 is oxidized by O released in reaction (16.1). The interaction of SO3 with H2O gas produces sulfuric acid, H2SO4·nH2O(l), the main component of cloud aerosols. Gravitational subsidence of aerosol particles causes evaporation of sulfuric acid and possible native sulfur to form H2SO4, H2O, and Sn gases at the cloud deck.  \n<!-- image -->  \nMetal (Fe, Mg, Ca, Ni, Na, Al) sulfates could precipitate at the cloud deck following the dissolution of space dust in the sulfuric acid aerosol and its evaporation (Zolotov et al. 2023). In the lower atmosphere, the thermal decomposition of ferric sulfate particles releases SO3 (or SO2 + O2) into the atmosphere (Mogul et al. 2024)  \nFe2(SO4)3 → Fe 2O3 (hematite) + 3SO3 (16.4)  \nfollowing by SO3 reduction to SO2. The chemical instability of settled ferric sulfate grains at the surface (Zolotov 2021) also suggests alteration to hematite and SO2 via interaction with reduced gases such as CO or OCS.  \nAs temperature increases below the cloud deck, SO3 forms through pyrolysis of H2SO4(g)  \nH2SO4 (g) → SO 3 + H2O. (16.5)  \nAt mid-30 km, thermochemical interactions of SO3 with OCS significantly consume both gases. However, reaction pathways are uncertain ( Sects. 16.3.1.2, 16.3.1.4, 16.3.1.6 ). The consumption of H2SO4 and SO3 is reflected in compositional profiles modeled in the sub-cloud atmosphere (Figs. 16.17). The photochemical production of CO above clouds via (16.1) and thermochemical consumption of OCS together with a poleward Hadley cell type circulation could cause decreasing x OCS and increasing x CO toward high latitudes (Fig. 16.16) observed in the sub-cloud atmosphere above 30 km ( Sect. 16.2.1.3 ). Photodissociation of CO2 (Eq. (16.1)), oxidation of SO2 to SO3 (Eq. (16.3)) and H2SO4 at cloud top, decomposition of H2SO4(g) to SO3 (Eq. (16.5)) and reduction of SO3 to SO2 by OCS and possibly CO below clouds, together with an upward transport of SO2 competes a short sulfur cycle (Fig. 16.15).  \nA long atmospheric cycle involves high-latitude subsidence of OCS-depleted and CO-enriched air and partial thermochemical conversion of CO and sulfur gases (S, S2) to OCS in the lowermost scale height (<16 km). OCS-rich near-surface gases are transported upward through eddy diffusion and equatorial rising, followed by the consumption of OCS in reactions with atomic S and SO3, as noted above. Estimated durations of short and long sulfur atmospheric cycles are months and 3 × 10 4 yr, respectively (Krasnopolsky 2013). Models for the current atmosphere require neither degassing from the interior ( Sect. 16.4.3 ) nor supply of gases through slow gas-solid interactions ( Sect. 16.3.2 ) that may not affect  \n<!-- image -->  \ncurrent atmospheric processes.  Sources and sinks of S-bearing gases are outlined in Table 16.6 and Table 16.7, which depict significant formation and consumption by gas-phase reactions.", '16.3.1.2  Sulfur Dioxide': 'Except for the upper mesosphere, SO2 is the most abundant S-bearing gas (Table 16.1). The low mixing ratio of SO2 in the upper and middle atmosphere reflects the net SO2 → H 2SO4 conversion in upper clouds. Atmospheric models suggest SO2 consumption in the middle atmosphere and production below clouds. Krasnopolsky (2013) inferred the SO2 lifetime of the order of months in the middle atmosphere. In the model of Bierson and Zhang (2020), the transport timescale of SO2 is ∼ 10 yr near the surface and months at 90 km. The chemical loss timescale is ~10 2 -10 9 yr in the lower and middle atmosphere and significantly lower above the upper middle atmosphere.', 'Consumption and Production at Cloud Top and Above': "In the middle atmosphere, SO2 is consumed through photodissociation that mainly occurs below 70 km altitude and oxidation to SO3 by (16.3), as inferred in many models (e.g., Yung and DeMore 1982; Krasnopolsky and Parshev 1983; Zhang et al. 2012; Krasnopolsky 2012, Fig. 16.18). Photodissociation is reversible, but the formation of H2SO4 results in the net loss of SO2. An additional net loss may occur if Sn gases and condensed Sx form in upper clouds (e.g., Pinto et al. 2021; Fig. 16.19) and sink with aerosol particles ( Sect. 16.3.1.7 ). Photolysis of SO2 leads to the formation of SO, (SO)2, and oxygen species, mainly O and O2, that are involved in recombination reactions to SO2 and in oxidation reactions to produce SO3 that hydrates to H2SO4 (Yung and DeMore 1982; Mills and Allen 2007; Krasnopolsky 2012; Zhang et al. 2012),  \nSO2 + hv → SO + O (16.6)  \nSO2 + O + M → SO 3 + M  \n(16.7)  \nSO3 + H2O + H2 O → H 2SO4(g) + H2O (16.8)  \nNet process: 2SO2 + H2 O → SO + H 2SO4(g)  \n(16.9)  \nA high H2SO4/SO concentration ratio at cloud tops (Table 16.1) suggests a significant contribution of an additional O in the SO2 → H 2SO4 conversion. The oxygen is supplied mainly through photolysis of CO2  \n47  \n<!-- image -->  \nabove clouds by (16.1) and drives the oxidation of S(IV) to S(VI) by (16.7). The primary net process of SO2 consumption and H2SO4(g) production in the middle atmosphere is  \nSO2 + CO2 + H2 O → H 2SO4(g) + CO. (16.10)  \nModeling and observations demonstrate that forming H2SO4(g) and liquid sulfuric acid (back reaction (16.5)) in upper clouds reduces x SO2 from ~130 -180 ppmv in the lower atmosphere to ~0.1 ppmv at ~70 km.  \nCl-bearing compounds (Cl, ClCO, ClO, etc.) play a crucial role in CO oxidation to CO2, which affects the net production of O, and they could be involved in SO2 oxidation (Yung and DeMore 1982; Mills and Allen 2007). DeMore et al. (1985) suggested an oxidation pathway that includes the interaction of SO2 with Cl, leading to the formation of ClSO2 ( Sect. 16.3.1.10 ), its oxidation to ClSO4 by O2, and the formation of SO3 through the interaction of ClSO4 with Cl. This pathway of SO2 consumption to H2SO4(g) corresponds to a net reaction  \nSO2 + Cl + O2 + H2 O → H 2SO4(g) + ClO. (16.11)  \nHowever, rate constants of corresponding reactions (e.g., ClSO2 oxidation) are uncertain, and subsequent models (Mills and Allen 2007; Krasnopolsky 2007, 2012, 2013; Bierson and Zhang 2020; Dai et al. 2024) do not support critical roles of Cl-bearing species in SO2 oxidation.  \nIn addition to photodissociation and oxidation, Krasnopolsky (2007) considered net SO2 reduction by CO  \nSO2 + 2CO → 2CO 2 + S (16.12)  \nthat could be a source of Sn in the clouds in O-deficient conditions. His models suggested exceeding H2SO4(g) flux (Eq. (16.10)) over that of Sn (Eq. (16.12)) by a factor of 4. An advanced model of Krasnopolsky (2013) demonstrates the overwhelming dominance of the net pathway (16.10) and suggests a supply of possible cloud sulfur species from the lower atmosphere rather than from the cloud top. The prevalence of reaction (16.10) agrees with the high H2SO4/Sx ratio (~7 -10) in cloud aerosols tentatively inferred from the Vega reaction gas chromatography data (Porshnev et al. 1987).  \nThe modeled SO2 consumption corresponds to its lifetime of the order of months in the middle atmosphere (Krasnopolsky 2007, 2013). A decrease in x SO2 above clouds toward higher latitudes ( Sect.  \n48  \n<!-- image -->  \n16.2.1.1 ) could reflect SO2 loss as gases move toward poles at the cloud's altitude (Marcq et al. 2013) with a typical SO2 lifetime of a few Earth days. The observed decrease in x SO2 above clouds at 70 -80 km (Fig. 16.18) is attributed to photochemical oxidation to SO by (16.6) (e.g., Zhang et al. 2012; Krasnopolsky 2012).  \nDespite the net SO2 loss through photodissociation and oxidation below ~80 km (Eqs. (16.6) -(16.12)), SO2 is regenerated through recombination of SO and O and via SO oxidation in reactions mostly with NO2 and ClO  \nSO + O + M → SO2 + M (16.13) SO + NO2 → SO2 + NO (16.14) SO + ClO → SO2 + Cl (16.15)  \n(Mills and Allen 2007; Krasnopolsky 2006, 2012; Zhang et al. 2010, 2012; Dai et al. 2024). Another likely source of SO2 is the oxidation of OCS ( Sect. 16.3.1.7 ) by the net process  \nOCS + 2CO2 → SO 2 + 3CO (16.16)  \nin the middle atmosphere (Krasnopolsky 2013) and above. Oxidation of H2S leads to SO2 as well. Another source of SO2 gas in the upper middle atmosphere and, possibly above, is SO2 exsolution from sulfuric acid aerosol (Rimmer et al. 2021; Dai et al. 2024).", 'Production and Consumption in the Lower Atmosphere': "Thermochemical reduction of SO3 by OCS and CO mainly accounts for the gain of SO2 in the lower atmosphere. It is unclear whether the reaction  \nSO3 + CO → CO 2 + SO2 (16.17)  \nis (Prinn 1975, 1985a; von Zahn et al. 1983; Bierson and Zhang 2020) or not (Krasnopolsky 2007, 2013) important below the cloud deck because its rate has not been determined experimentally ( Sect.  \n16.3.1.4 ). Krasnopolsky and Pollack (1994) and Krasnopolsky (2007, 2013) modeled SO2 production through a coupled reduction of SO3 and oxidation of OCS by net process  \n49  \n<!-- image -->  \nSO3 + 2OCS → CO 2 + CO + SO2 + S2 (16.18)  \nthat may include reactions  \nOCS + SO3 → CO 2 + (SO)2 (16.19)  \n(SO)2 + OCS → CO + SO 2 + S2 (16.20)  \nTheir thermochemical kinetic and eddy diffusion models suggest a strong consumption of SO3 at ~36 km altitude (Fig. 16.17). The validity of this pathway remains to be tested through experimental assessments of reaction rates and detection of short-living (SO)2 at 30 -40 km. Bierson and Zhang (2020) included the fast rate of (SO)2 → 2SO conversion in their model and concluded that (SO) 2 may not build up to make reaction (16.20) efficient. To avoid this problem, they modeled SO3 loss via three-body process  \n2SO3 + OCS → 3SO 2 + CO, (16.21)  \nand reproduced the observed OCS gradients at 33 -36 km ( Sect. 16.2.1.3) and inferred a positive CO gradient with altitude, roughly consistent with observations. Dai et al. (2024) adopted (16.21) in their models and observed a match with x OCS data at 30s km. Interactions of OCS and SO3 and formation of SO2 and CO are consistent with observed anticorrelation of CO and OCS and vertical and latitudinal gradients of OCS, CO, and SO2 ( Sect. 16.2.1 ). The accuracy of assessing contributions from reactions (16.19), (16.20), and (16.21) is limited due to the absence of experimental data on reaction rates. However, regardless of SO3 reduction reactions, all considered pathways suggest a source of SO2 at ~35 -40 km. From the cloud deck to lower 30s km, some SO2 could also be produced by reducing H2SO4(g) by CO (Mills et al. 2007) and OCS. All considered interactions could be more important in high latitude downwelling of Hadley cell circulation (Fig. 16.16), and the elevated x SO2 at 33 km and higher latitudes (Marcq et al. 2021) could reflect SO2 production via reduction of H2SO4(g) and SO3. If so, no SO2 consumption in the deep atmosphere must be invoked to explain observations.  \nDeep atmospheric SO2 could be produced and consumed in thermochemical reactions considered in the kinetic models of Krasnopolsky and Pollack (1994), Krasnopolsky (2007, 2013), Bierson and Zhang (2020), and Dai et al. (2024). Although low x SO2 values in the low latitude cloud (~51 -54 km, Oschlisniok et al. 2021) and sub-cloud atmosphere (~33 km, Marcq et al. 2021) ( Sect. 16.2.1.1 ) suggest equatorial  \n<!-- image -->  \nupwelling of SO2-depleted air, no major SO2 sink in the lower atmosphere has been modeled.  \nThermochemical reactions that consume SO2 are likely compensated by back reactions (Krasnopolsky 2007, 2013) and could reach thermochemical equilibria at the surface (Krasnopolsky and Parshev 1979; Krasnopolsky and Pollack 1994; Zolotov 1996; Fegley et al. 1997b, Table 16.2). A slow geological sink of SO2 via gas-solid type reactions (Fig. 16.15, Sect. 16.3.2.2 ) may not be responsible for the apparent SO2 deficiency in the deep low-latitude atmosphere. Net reduction of SO2 by CO (Eq. (16.12)) in the lower scale height by reactions  \nSO2 + CO → CO 2 + SO (16.22)  \nSO + SO → SO 2 + S (16.23)  \nmainly contribute to SO2 loss in Krasnopolsky's (2013) and Dai et al. (2024) models. The interaction of two S2O mainly compensates for SO2 loss to form SO2 and S3 Dai et al. (2024) model.  \nConsumption of SO2 by a net thermochemical process  \nSO2 + 3CO → OCS + 2CO 2 (16.24)  \nmay not be an essential pathway in the lower atmosphere based on kinetic modeling (Krasnopolsky 2007, 2013; Bierson and Zhang 2020; Dai et al. 2024) and experimental studies at Venus' temperatures and 1 bar (Hong and Fegley 1997a). However, the stoichiometry of (16.24) does not contradict a loss of 20 -30 ppmv SO2 (Marcq et al. 2021) and a gain of 20 -30 ppmv OCS (Pollack et al. 1993; Krasnopolsky and Pollack 1993; Krasnopolsky 2007, 2013; Fegley et al. 1997b) in the deep low latitude atmosphere. Possible chemical equilibration between SO2, OCS, CO, and CO2 only at the conditions of modal planetary radius (Krasnopolsky and Pollack 1994; Zolotov 1996; Fegley et al. 1997b) may reflect catalytic effects of surface materials (e.g., iron oxides and sulfides, Sect. 16.3.2.3 ) that allow interaction of absorbed gas reactants and accelerate the reduction of SO2. The uncertainty surrounding the potential depletion of SO2 in the deep lower atmosphere will be resolved with the analysis of gas abundances conducted by the DAVINCI descent sphere (Garvin et al. 2022, Sect. 16.5 ).", 'Causes of Spatial and Short- and Long-term Variability': "In the mesosphere and at the cloud top, both spatial and temporal (over a scale of a few days) variations in x SO2 within two or more orders of magnitude (Fig. 16.2, Sandor et al. 2010; Belyaev et al. 2012, 2017;  \n<!-- image -->  \nEncrenaz et al. 2012, 2016, 2019, 2020; Vandaele et al. 2017a, 2017b; Marcq et al. 2013, 2020; Mahieux et al. 2023) likely reflect a short photochemical life of SO2 rather than winds (e.g., Encrenaz et al. 2012). This is consistent with a daily cloud top SO2 cycle (Marcq et al. 2020) and higher x SO2 and SO2/SO ratios observed at the night side at 85 and 95 km (Sandor et al. 2010; Belyaev et al. 2017). The photochemicaldynamic model of Shao et al. (2022) shows that a sizable day-night difference in x SO2 above 85 km results from both photochemistry and the subsolar-to-antisolar circulation. In general, the effects of winds and circulation on the variability decrease with altitude. An irregular flux of cosmic materials (Carrillo-Sánchez et al. 2022, Table 16.12) may also affect S-bearing gases in the mesosphere.  \nAt the cloud top, SO2 variability reflects competition between photochemical destruction and supply via advection, which is currently higher at lower latitudes (Encrenaz et al. 2019; Marcq et al. 2020).  A correlation of transient SO2 enhancement at cloud tops with local low latitude upwellings (plumes, 3 m s -2 measured by Vega balloons, Linkin et al. 1986) agrees with supply from the sub-cloud atmosphere through general meridional circulation (Encrenaz et al. 2019; Jessup et al. 2020; Marcq et al. 2013, 2020). Marcq et al. (2013) showed that even a tiny (~1%) variation in SO2 cloud supply through low latitude upwelling could cause tens of ppbv latitudinal variations in mesospheric SO2. Marcq et al. (2020) noted that the decrease in x SO2 around local solar noon at low latitudes indicates that local SO2-rich air plumes cannot counterbalance the rapid mid-day photochemical depletion. The links between heat balance, momentum, convection, and turbulent mixing in clouds were first inferred in three-dimensional models of Lefèvre et al. (2018, 2022). They reproduced a diurnal cycle in cloud convection and predicted a 7-km thick convective layer at the cloud top caused by absorption of the solar UV light. SO2, as a chemical tracer, was included by Leférve et al. (2022) to assess the effect of cloud dynamics on spatial and temporal variability. The estimated vertical eddy diffusion aligned with assessments from in situ measurements, but it appeared to be several orders of magnitude compared to the values used in 1D chemical modeling. The models of Kopparla et al. (2019) showed that observed chemical variability patterns could be related to the 4-day Kelvin wave, 5-day Rossby waves, and the overturning circulation. The observations of Jessup et al. (2015) and Marcq et al. (2020) suggest that the circulation regime with many low-latitude plumes and a decrease in x SO2 toward higher latitudes (Fig. 16.2) could be replaced within a few Earth days by a regime with fewer plumes and a reversed latitudinal gradient.  \nAn enrichment in cloud top SO2 and an enhanced UV brightness above the downwind western slopes of Aphrodite Terra (Bertaux et al. 2016; Jessup et al. 2015, 2020; Marcq et al. 2020) may indicate a vertical  \n<!-- image -->  \ntopography-related supply mechanism, possibly related to the phase-shifted vertical winds (Bertaux et al. 2016). In turn, the typography-related winds are likely linked to stationary gravity atmospheric waves that correlate with topography (Peralta et al. 2017; Fukuhara et al. 2017; Koyama et al. 2019; Kitahara et al. 2019; Lefevre et al. 2020; Suzuki et al. 2023). Mountain gravity waves were suggested from the UV and mid-IR absorption patterns at the cloud top observed by Akatsuki (Fakuhara et al. 2017; Peralta et al. 2017). Kouyama et al. (2017) emphasized the impact of solar illumination, local time, and latitude on the waves' appearance in Akatsuki cloud top mid-IR data. The models of Leférve et al. (2018) revealed the effects of gravity waves on convection and mixing in clouds. Overall, the physical processes in the lower atmosphere could influence the dynamics (plumes and turbulent mixing) and short-term chemical variability throughout clouds, as shown in the models of Moriella et al. (2022), Leférve et al. (2018, 2022) and Kopparla et al. (2019). In lower clouds, the compositional variability of SO2 and H2SO4(g) (Fig. 16.5) is mainly affected by turmoil in atmospheric circulation, while some SO2 latitudinal dependence suggests SO2-depleted plumes at low latitudes (Oschlisniok et al. 2021).  \nThe decline of the disc-averaged cloud top x SO2 from ~500 ppbv in 1978 to ~20 ppbv in 1995 and a drop from ~200 ppbv in 2007 to ~10 ppbv in 2014 (Fig. 16.3, Esposito et al. 1988; Encrenaz et al. 2013, 2016; Marcq et al. 2013, 2020; Vandaele et al. 2017b) has been a discussion topic for 40 years. These changes have been initially interpreted as volcanic eruptions providing a buoyant plume that lifts SO2 into the visible atmosphere and an uneven supply of SO2 through volcanic degassing (Esposito 1984; Esposito et al. 1988). However, this is uncertain because of the ambiguity of current volcanism and corresponding pyroclastic activity (Herrick et al. 2023, for review) and suppressed magma degassing at ambient pressure of 10 2 bars ( Sect. 16.4.3 ).  Other models explained the viability by periodic changes in effective eddy diffusion in clouds (Krasnopolsky 1986) and in global circulation (Clancy and Muhleman 1991). Kouyama et al. (2019) and Kitahara et al. (2019) suggested that the variability reflects momentum deposition from propagating atmospheric gravity waves induced by topography. Marcq et al. (2020) noted that periods of high SO2 abundance in the early 1980x and late 2000s could have multiple lowlatitude plumes. In other words, changes in cloud circulation regime, mainly in low-altitude plumes, could cause both temporal variability and long-term trends observed at the cloud tops since the 1960s.  \nThe long-term anti-correlation of SO2 and water vapor at cloud tops observed in 2014 -2019 by Encrenaz et al. (2020) but not in 2021 (Encrenaz et al. 2023) may indicate H2O involvement in processes that affect SO2. A one-dimensional photochemical-diffusion cloud model of Shao et al. (2020) demonstrates  \n<!-- image -->  \nclose inter-relations of SO2 and H2O in the middle and upper clouds (~58 -65 km) and provides results consistent with the anti-correlation. The modeling shows that eddy mixing transport alone cannot explain the observations and suggests a significant influence of sub-cloud processes. The models of Kopparla et al. (2020) demonstrate that long-term changes in cloud convection patterns could be caused by convective strength oscillations linked to the radiative effects of water vapor abundance at the cloud base. In their model, x H2O at the convective cloud deck affects clouds' thermal balance and convection, and vice versa. Kopparla et al. (2020) noted that the oscillation timescale of 3 -9 years inferred in the model reflects the geometric mean of the radiative cooling time and the eddy mixing time at the deck of convective clouds. Shao et al. (2020) modeled the effects of periodic changes in chemical processes in upper clouds and demonstrated that water vapor supply to the clouds via non-diffusive mechanisms (e.g., plumes) could affect SO2 at cloud tops and a long-term anti-correlation between SO2 and H2O observed at 64 km by Encrenaz et al. (2020). Rimmer et al. (2021) considered the effects of periodic long-term depletions in sub-cloud SO2 and hypothesized exotic mechanisms that could be responsible for such depletions.", 'SO2 Gradients': 'The cause of increasing x SO2 with altitude at ~75 -100 km ( Sect. 16.2.1.1 , Fig. 16.18) is not fully understood and conflicts with predictions from photochemical models of Yung and DeMore (1982) and Krasnopolsky (2012). Several researchers suggested a contribution from an unknown sulfur reservoir that allows the production of SO2 and SO through upper mesospheric photochemistry (Sandor et al. 2010, 2012; Belyaev et al. 2012, 2017; Vandaele et al. 2017a, 2017b). Sandor et al. (2010) first discussed a supply of SO2 from a mesospheric aerosol (Wilquet et al. 2009) that could consist of condensed sulfuric acid. The photochemical models by Zhang et al. (2010, 2012) demonstrated that the observed profiles of SO₂ and SO could be replicated by a significant increase in the concentration of H₂SO₄(g). However, the assessed observational upper limit for x H2SO4(g) above 85 km (Table 16.1) ruled out the H2SO4(g) as a source of sulfur excess. Belyaev et al. (2012) and Mahieux et al. (2024) pointed out that a warmer upper mesosphere promotes evaporation of sulfuric acid aerosol if it exists in that region. Belyaev et al. (2017) noted that the SO2 enrichment above 85 km correlates with the density of putative supersaturated sulfuric acid droplets in the upper haze at 70 -90 km, as assessed by Luginin et al. (2016). Model sensitivity studies of Parkinson et al. (2015) demonstrated a possibility of increasing xSO2 above 80 km at specific eddy diffusion profiles and SO2 and water vapor contents in upper clouds. Sandor et al. (2012)  \n<!-- image -->  \nproposed elemental sulfur as a probable reservoir in the upper mesosphere required by observed SO2 and SO abundances to conserve S atoms. Chemical and photochemical models of Zhang et al. (2012) and Zhang and Bierson (2020) reproduced the observed SO2 and SO contents above 75 km by forcefully invoking a delivery of condensed polysulfur (Sx, S8) to the mesosphere from above. However, no observation supports the existence of an Sx aerosol there. The nominal model of Dai et al. (2024) did not reveal elevated x SO2 above 80 km, and the authors suggested a non-photochemical sulfur source. Vandaele et al. (2017a) stated that no current photochemical model reproduces the SO2 profile above ~75 km without significant manipulation of the physical properties of the atmosphere. It remains to be estimated whether the daily delivery of 1 -4 metric tons of cosmic sulfur (Table 16.12) could explain the mesospheric profile of SO2 and enhanced abundances of SO, SO3, OCS, CS, and CS2 (Table 16.1).  \nThe latitudinal gradients of SO2 at and above the cloud top (Marcq et al. 2011, 2013, 2020; Belyaev et al. 2012; Encrenaz et al. 2019; Jessup et al. 2015; Fig. 16.2), in the middle cloud (Oschlisniok et al. 2021), and in the lower atmosphere at 33 km (Encrenaz et al. 2020; Marcq et al. 2021, Fig. 16.4) could reflect the fate of SO2 in its transport through the Hadley cells (Fig. 16.16) that may involve the mesosphere. The often observed decrease in cloud-top SO2 toward poles could reflect consumption (Eqs. (16.6), (16.7)) and lesser plume activity at higher latitudes (Encrenaz et al. 2019; Marcq et al. 2020). In the mesosphere, the latitudinal gradient could manifest SO2 propagating to the upper mesosphere at low latitudes and circulation poleward followed by a downward to the lower mesosphere (Vandaele et al. 2017a).  \nThe enhanced high latitude x SO2 at ~33 km (Table 16.1) could reflect SO2 formation through the reduction of SO3 (Eqs. (16.18) -(16.21)) and H2SO4(g) in a sub-cloud downwelling (Fig. 16.16). Alternatively, or additionally, lower x SO2 at low latitudes at ~33 km (Marcq et al. 2021) and at 51 -54 km (Oschlisniok et al. 2021) does not exclude some consumption of SO2 in the deep atmosphere followed by equatorial upwelling of an SO2-depleted gas, as discussed above. Such an upwelling is roughly consistent with the interpretation of in situ UV spectra obtained at Vega probes at 10 -60 km, suggesting a decrease in x SO2 toward the surface (Bertaux et al. 1996). However, lower atmosphere models (Krasnopolsky and Pollack 1994; Krasnopolsky 2013, 2013; Bierson and Zhang 2020; Dai et al. 2024) do not predict any SO2 vertical gradient. As mentioned earlier, the effectiveness of SO2 reduction to OCS (Eq. (16.24)) in the near-surface atmosphere could be limited. It is unclear what may cause a significant (tens of ppmv) net consumption of SO2 in the deep equatorial atmosphere if it occurs.', '16.3.1.3  Sulfur Monoxide and Disulfur Dioxide': 'Although SO is observed only in the middle and upper atmosphere ( Sect. 16.2.1.1 ; Table 16.1), its abundance and reaction pathways have been evaluated in photochemical and thermochemical kinetic models throughout the atmosphere.  None of the three disulfur dioxide dimers ((SO2)2, cis, trigonal, and trans) are detected, but kinetic modeling suggests that dimers are important intermediate species. At cloud tops and in the mesosphere, SO is the main product of SO2 photolysis (16.6). The formation of sulfuric acid is ineffective above ~75 km. SO is the main photochemical product of SO2 in the mesosphere. An increase in x SO2 and x SO above 80 km suggests a common source of sulfur (Sandor et al. 2010, 2012; Belyaev et al. 2012) ( Sect. 16.3.1.2 ). A more efficient SO2 photolysis in the upper mesosphere at 90 -100 km likely causes elevated dayside SO/SO2 mixing ratios there (Belyaev et al. 2012, 2017). On the night side, the suppressed photolysis accounts for lower x SO and SO/SO2 ratios (Sandor et al. 2010). The only partial short-term, long-term, and spatial correlations of SO and SO2 at cloud tops and in the mesosphere (Sandor et al. 2012; Belyaev et al. 2012, 2017; Jessup et al. 2015; Encrenaz et al. 2015) suggest that SOx photolysis is not the only source for sulfur in the upper atmosphere, as first noted by Jessup et al. (2015).  \nThe high compositional and spatial variability of SO (Sandor et al. 2010, 2012; Belyaev et al. 2012;  \nEncrenaz et al.  2015; Jessup et al. 2015) reflects its short chemical lifetime (hours or less, Bierson and Zhang 2020, Dai et al. 2024) and reasons that affect SO2 fluctuations in and above clouds ( Sect. 16.3.1.2 ). These inferences have been assessed in photochemical models of Yung and DeMore (1982), Mills and Allen (2007), Yung et al. (2009), Zhang et al. (2012), Krasnopolsky (2012), Zhang and Bierson (2020), and Dai et al. (2024) in which SO is involved in multiple reactions. Although SO is consumed via photolysis to atomic S and O, it mainly converts back to SO2 (Eqs. (16.13) -(16.15)). In the model of Mills and Allen (2007), reaction (16.15) provides 10 -20% SO loss at 66 -80 km. The model of Krasnopolsky (2012) suggests the SO + NO2 interaction (16.14) as the most likely SO loss below 75 km. That model also implies conversion SO to SO2 via intermediate (SO)2 by reactions  \nSO + SO + M → (SO)2 + M (16.25)  \n(SO)2 + hv → SO2 + S (16.26)  \nas the main pathway of atomic S production in the middle atmosphere ( Sect. 16.3.1.7 ).  \nThe model of Dai et al. (2024) suggests a major SO sink by (16.25) at 60 -70 km, a significant consumption by (16.14) at 75 -80 km, and losses by (16.13) and photolysis at 80 -112 km  \nSO + hv → S + O (16.27)  \nTheir nominal model revealed the observed increase in x SO in the upper mesosphere, suggesting increasing production with altitude above 80 km via reaction  \nS + O2 → SO + O (16.28)  \nthat provides the major SO source above ~105 km. Dai et al. (2024) modeled that a slow release of dissolved SO2 from sulfuric acid aerosols contributes to the supply of mesospheric SO2 to the increase in x SO with altitude.  \nSignificant attention has been devoted to SO dimers in upper clouds because of their decent spectral match with the spectra of clouds (Frandsen et al. 2016; Pérez-Hoyos et al. 2018) ( Sect. 16.3.1.8 ). However, (SO)2 species have too short photochemical life (2 to 5 s) to be abundant and contribute to the blue and UV parts of spectra (Frandsen et al. 2016, 2020). At 60-70 km, the model of Dai et al. (2024) suggests SO regulated by rapid conversions with (SO)2 ((16.15) and back reaction). Neither the photochemical models of (Krasnopolsky 2018; Bierson and Zhang 2020; Dai et al. 2024) nor observations (Marcq et al. 2020) indicate a sufficient SO to form an abundant (SO)2 in the middle atmosphere. In contrast to other models, Rimmer et al. (2021) assessed increasing x (SO)2 from 80 km to the surface. The role of (SO)2 in the formation of Sn and S2O in upper clouds is emphasized in the photochemical models of Pinto et al. (2021) and Francés-Monerris et al. (2022) ( Sect. 16.3.1.7, 16.3.1.9 ).  \nIn the lower atmosphere, SO could be an essential intermediate species in thermochemical reactions, as modeled by Krasnopolsky and Pollack (1994) and Krasnopolsky (2007, 2013). Although (SO)2 has been considered as an intermediate species in OCS + SO3 interactions at ~36 km ((16.18) -(16.20); Krasnopolsky and Pollack 1994; Krasnopolsky 2007, 2013), fast thermochemical consumption of (SO)2 and SO (Mills 1998) could impede the role of (SO)2 in OCS oxidation, as noted by Bierson and Zhang (2020). In the model of Krasnopolsky (2013), SO forms through SO2 + CO interaction (16.22) and reaction  \nSO2 + S → SO + SO (16.29)  \n<!-- image -->  \nwith maximal yields at 15 km and 5 km, respectively. SO is consumed via a slow SO + SO interaction (16.23), which could be a major net source of S gas in the lowest scale height. Below 40 km, the model of Dai et al. (2024) suggests SO production via (16.22), (16.29), and back reaction (16.25). Despite the importance of SO in thermochemical reactions, it is not a measurable species with a modeled mixing ratio of 0.1 -13 pptv below 40 km (Krasnopolsky et al. 2007).  At the surface conditions of modal planetary radius, chemical equilibrium x SO is 17 -41 pptv (Table 16.2).', '16.3.1.4  Sulfur Trioxide': 'Although SO3 is only detected in the mesosphere above 75 km (Mahieux et al. 2023), it is an essential reactant above ~35 km, as inferred in photochemical and thermochemical models. Surprisingly, SO3 is the most abundant S-bearing gas in the upper mesosphere (Table 16.1). The reported x SO3 levels above 75 km are 1 -3 orders of magnitude higher than those modeled by Zhang et al. (2012), Krasnopolsky (2012), Bierson and Zhang (2020), and Dai et al. (2024), which were based on measured x SO2 and the upper limit of H2SO4(g).  The observed increase in x SO3 with altitude disagrees with modeling (Krasnopolsky 2012) performed without a forceful addition of S-bearing species. A sensitivity modeling of Bierson and Zhang (2020) shows that x SO3 is weakly affected by mesospheric x H2SO4 but strongly depends on an additional source of native sulfur (Sx, S8) that also may cause gradients of SO2 and SO above 80 km ( Sects. 16.3.1.2, 16.3.1.3 ). The inconsistency between the observed and modeled SO3 remains to be resolved. The supposed delivery of condensed sulfuric acid to the mesosphere (e.g., Zhang et al. 2010; Belyaev et al. 2012, 2017; Mahieux et al. 2024) could be a source of SO3 via thermal evaporation and photochemical decay of H2SO4 gas and aerosol (Mahieux et al. 2023).  In addition, photochemical oxidation of cosmic S-bearing materials (mainly FeS, S-bearing organic matter, CarrilloSánchez et al. 2022) may also provide a reasonable explanation for positive gradients of mesospheric sulfur oxides.  \nIn upper clouds, SO3 is mainly produced through the oxidation of SO2 by atomic oxygen (Eq. (16.7)) formed via photolysis of CO2 by (16.1) and SO2 by (16.6). In upper clouds, SO3 is hydrated to H2SO4(g) by (16.8) and converted to liquid sulfuric acid,  \nH2SO4(g) + nH2 O(g) → H 2SO4·nH2O(l). (16.30)  \n<!-- image -->  \nSome SO3 is consumed via photolysis to SO2 and O and via reactions with SO and O (Zhang et al. 2012; Krasnopolsky 2012). The formation of sulfuric acid significantly reduces both x SO2 and xSO3, and SO3 is modeled as a trace gas throughout clouds with a maximal mixing ratio of ~1 ppbv at ~67 km (Zhang et al. 2012; Krasnopolsky 2012). Below clouds, SO3 forms via thermal dissociation of H2SO4(g) by reaction (16.5), as inferred in kinetic models (Krasnopolsky and Pollack 1994; Krasnopolsky 2007, 2013; Bierson and Zhang 2020). These models suggest xSO3 increasing with lowering altitude from the cloud deck to ~38 km (at low latitudes), where xSO3 reaches ~0.4 ppmv (Fig. 16.17). The decrease in x H2SO4(g) at lower altitudes observed by Kolodner and Steffes 1998 (Fig. 16.5) and Imamura et al. (2017) at 35 -47 km likely reflects the thermal dissociation of the gas.  \nThermochemical kinetic models for the sub-cloud atmosphere imply SO3 consumption through reactions with OCS and CO below the cloud deck down to ~30 km. The models of Krasnopolsky and Pollack (1994), Krasnopolsky (2007, 2013), Bierson and Zhang (2020), and Dai et al. (2024) suggest a significant SO3 consumption at ~36 through thermochemical interactions with OCS (Eqs. (16.18) -(16.21)). Although the validity of the modeled SO3 consumption determined by reactions with OCS supports the observed OCS gradient at 30 -48 km (Table 16.1), actual reaction pathways are uncertain because of insufficient experimental data on reaction rates. Krasnopolsky et al. (2007, 2013) concluded that the SO3-CO interaction (16.17), previously considered by Prinn (1985a) and Krasnopolsky and Pollack (1994), provides only a minor (~4 %) contribution to the loss of sub-cloud SO3. A subordinate role of reaction (16.17) agrees with a consistent x CO at 30 -40 km and a substantial decrease in x OCS with altitude that could be caused by interactions of OCS with SO3 (Eqs. (16.18) -(16.21)). However, Bierson and Zhang (2020) showed that reaction (16.17) could be critical at 30 -50 km, especially at higher altitudes where SO3 is produced from H2SO4(g) by (16.5). The participation of CO in SO3 reduction does not contradict the higher x CO compared to x OCS at 33 -48 km (Table 16.1). There is no observational data on CO at 42 -50 km to verify the role of reaction (16.17), though modeling of Bierson and Zhang (2020) shows that a decrease in x CO may not occur at those altitudes. The consumption of CO by reaction (16.17 could be compensated by CO production via interactions of SO3 with OCS (Eqs. (16.18) -(16.21)). The modelers have relied on the reaction (16.17) rate, as estimated by Krasnopolsky and Pollack (1994). However, the precise contribution of reaction (16.17) to the consumption of SO3 remains uncertain until its rate is experimentally determined. Despite the uncertain roles of CO and OCS, all models suggest an  \n<!-- image -->  \ninsignificant role of SO3 in atmospheric chemistry below ~30 km, and thermochemical equilibrium models for the near-surface atmosphere suggest only 0.3 -0.9 pptv SO3 (Table 16.2).', '16.3.1.5  Sulfuric Acid Gas and Aerosol': "Sulfuric acid vapor has been detected in lower clouds and below the cloud base at 30 -55 km ( Sect. 16.2.1.2 , Table 16.1). Sulfuric acid is the second major S-bearing gas within clouds down to ~35 -40 km. Its fate is linked to SO3, H2O, and sulfuric acid aerosol (Fig. 16.15), as evidenced by inferences made in photochemical and thermochemical models (Yung and DeMore 1982; Krasnopolsky and Parshev 1981, 1983; Krasnopolsky and Pollack 1994; Mills 1998; Mills and Allen 2007; Yung et al. 2009; Zhang et al. 2012; Krasnopolsky 2007, 2012, 2013; Bierson and Zhang 2020). H2SO4(g) forms in upper clouds via the oxidation of SO2 to SO3 by (16.7) followed by hydrolysis of SO3 and H2SO4(g) (Eqs. (16.8), (16.30)), leading to the formation of liquid sulfuric acid, H2SO4·nH2O(l), where n is 1.53 -0.47 in 78 -92 % sulfuric acid by mass (Titov et al. 2018), and more concentrated acid is modeled in the lower cloud by Dai et al. (2022) and Shao et al. (2024). H2SO4 vapor coexists with sulfuric acid aerosol and forms at the cloud base via thermal evaporation of the aerosol (back reaction (16.30)). It thermally decomposes to SO3 and H2O below the cloud deck by (16.5) (Figs. 16.15, 16.16). The temperature at the bottom of the primary cloud at ~48 km (at low latitudes) agrees with the vaporization of liquid sulfuric acid aerosol particles (e.g., Krasnopolsky and Pollack 1994; Krasnopolsky 2015; Dai et al. 2022; Shao et al. 2024).  \nThe suggested photodissociation of mesospheric H2SO4(g) (Sandor et al. 2010) to account for the positive SO2 and SO gradients above 85 km ( Sect. 16.2.1.1 ) is inconsistent with the H2SO4(g) lower limit at 85 -100 km (Sandor et al. 2012). However, photochemical modeling shows that the observed profiles could be reproduced by adding H2SO4(g) produced via evaporation of sulfuric acid aerosol (Zhang et al. 2010, 2012; Bierson and Zhang 2020). Belyaev et al. (2012, 2017) and Mahieux et al. (2024) noted that elevated temperature at 90 -110 km favors the release of H2SO4(g) from aerosol particles that could exist in the mesosphere (Luginin et al. 2016). According to Mahieux et al. (2023), the frequency pattern observed for SO3 in the mesosphere indicates a conversion between SO3 and H2SO4(g). This conversion may be linked to condensed sulfuric acid, a potential source of H2SO4(g). The increased mixing rations of H2O and HDO in the upper mesosphere reported by Mahieux et al. (2024) was interpreted in terms of the thermal dissociation of sulfuric acid aerosols delivered from beneath.  \n<!-- image -->  \nProduction of H2SO4 vapor and sulfuric acid aerosol in upper clouds depends on the supply of SO2 from the lower atmosphere via eddy diffusion and plumes. Models that considered eddy diffusion (Krasnopolsky and Pollack 1994; Mills 1998; Zhang et al. 2012; Krasnopolsky 2012, 2015; Bierson and Zhang 2020) suggest the formation of H2SO4 in a narrow layer in upper clouds. As such, Krasnopolsky's (2012) model for low latitudes indicates the formation of H2SO4 in a 3 km thick layer with a peak at 66 km located just below a layer of the most productive SO2 photolysis. The low latitude upwelling likely boosts H2SO4 production at higher altitudes than middle latitudes (Figs. 16.2, 16.16). Photochemical modeling with circulation by Stolzenbach et al. (2023) and the 3D cloud model of Shao et al. (2924) suggest the production of H2SO4 from SO2 at higher altitudes in the equatorial upwelling region.  \nWithin the main cloud layer, measured x H2SO4(g) generally corresponds to the vapor pressure in equilibrium with sulfuric acid (Kolodner and Steffes 1998; Imamura et al. 2017; Oschlisniok et al. 2021; Dai et al. 2022; Shao et at. 2024) (Fig. 16.5). The H2SO4(g) mixing ratio increases with the temperature toward the cloud base. According to the models of Hansen and Hovenier (1974) and Pollack et al. (1978), the concentration of H2SO4 in the sulfuric acid aerosol increases from 75 -85% at ~68 km toward the cloud base, where it reaches ~98% by mass. In the models of Dai et al. (2022) and Shao et al. (2024), the acidity changes from 78 wt% H2SO4 at 62 km to 98% at the cloud base. Aerosol particles descend towards the cloud deck and evaporate, forming H2SO4(g) and water vapor (back reaction (16.30)). Fate of liquid sulfuric acid in clouds is consistent with the near constant x H2O (30 -50 ppmv) in clouds inferred from ground-based nightside near-IR observations (Pollack et al. 1993) and Venera 11, 13, and 14 in situ near-IR spectroscopy data (Ignatiev et al. 1997). In the model of Krasnopolsky and Pollack (1994), sulfuric acid aerosol evaporates at 48.4 km, consistent with in situ data from Pioneer Venus and Venera probes for the low latitude atmosphere.  In the improved cloud model for the H₂SO₄ -H₂O syst em, Krasnopolsky (2015) determined cloud base altitudes at different latitudes: 47.5 km globally, 48.5 km at low latitudes, and 46 km at 60° latitude. He assessed the concentrations of water vapor, H₂SO₄(g), and sulfuric acid within clouds. Several other models (e.g., Zhang et al. 2012; Krasnopolsky et al. 2012; Dai et al. 2022) predicted aerosol evaporation at ~47 -48 km.  \nHigher H2SO4(g) abundances and high variability observed at equatorial and polar latitudes below 55 km ( Sect. 16.2.1.2 , Table 16.1) could reflect atmospheric dynamics that influence temperature at certain altitudes and conditions of aerosol evaporation ( Sect. 16.3.1.2 ). The Hadley-type hot equatorial upwelling (Fig. 16.16) is proposed as a potential explanation for the occurrence of a sub-cloud H2SO4-rich  \n<!-- image -->  \nlayer at approximately 4 km higher altitude at equatorial latitudes than at polar latitudes (Oschlisniok et al. 2021). The lower and less variable x H2SO4(g) observed below 55 km in middle latitudes could be attributed to downwelling near the boundary of a Hadley circulation cell and a smaller polar circulation cell, as modeled by Oschlisniok et al. (2021). The observed accumulation of H2SO4(g) released from sulfuric acid aerosol at equatorial and polar latitudes has been reproduced by upward winds in their models. Dai et al. (2023) used Venus Express data on H2SO4(g) abundance to estimate vertical eddy diffusion from the cloud base (43 -48 km) to 55 km at different latitudes. Their approach included condensation of H2SO4(g) and diffusion. The assessed eddy diffusion coefficient is an order of magnitude larger than other observation-based and model results. Significant latitudinal variations in eddy diffusion are consistent with global circulation. However, the elevated mixing in clouds has little effect on atmospheric processes in the upper and lower atmosphere (Dai et al. 2023, 2024). The 3D model of Shao et al. (2024) includes global circulation, transport within and around clouds, and chemical processes in the H2SO4-H2O system at 40 -85 km, providing sufficient agreement with measured x H2SO4(g), x H2O, and aerosol properties at different latitudes and altitudes.", '16.3.1.6  Carbonyl Sulfide': 'OCS is the most abundant reduced S-bearing gas observed at the altitude of ~30 km to 100 km ( Sect. 16.2.1.3 ). The Fate of OCS on Venus has been considered and numerically modeled by Prinn (1975, 1978, 1985a, b), von Zahn et al. (1983), Yung and DeMore (1982), Krasnopolsky and Parshev (1981a, 1981b, 1981c, 1983), Krasnopolsky and Pollack (1994), Mills (1998), Krasnopolsky (2007, 2012, 2013), Yung et al. (2009), Mills and Allen (2007), Zhang et al. (2012), Bierson and Zhang (2020), and Stolzenbach et al. (2023). OCS forms in the deep lower atmosphere, undergoes transport towards the clouds through eddy diffusion and a low latitude upwelling, experiences thermochemical oxidation below clouds, and is depleted by photolysis in and above upper clouds. In the sub-cloud atmosphere, some OCS could be transferred to the near-surface atmosphere through global circulation subsidence at higher latitudes (Fig. 16.1). This chemical and circulation pattern is supported by the increased OCS contents at lower altitudes and the elevated x OCS at low latitudes above and below clouds.', 'Formation of OCS': "<!-- image -->  \nThe inverse relationship between the abundances of OCS and CO throughout the sub-cloud atmosphere, along with the increased x OCS/ x CO ratio at low latitudes ( Sect. 16.2.1.3 ), indicates OCS formation in a deep lower atmosphere, as discussed by Marcq et al. (2005, 2008, 2018, 2020) and modeled by Yung et al. (2009) and Marcq and Lebonnois (2013). Models of Pollack and Krasnopolsky (1994) and Krasnopolsky (2007, 2013) show that OCS production occurs at the lowest scale height (< 16 km). They show that OCS likely forms through net CO + Sn interaction, in which Sn stands for atomic S and S2, the most abundant sulfur gas in the near-surface atmosphere ( Sect. 16.3.1.7 , Fig. 16.17, Table 16.2). The models of Krasnopolsky (2007, 2013) suggest OCS formation primarily by reaction  \nS + CO + M → OCS + M (16.31)  \nthat exhibits the highest yield at an altitude of 9 km. The results imply that OCS production by reaction  \nS2 + CO→ OCS + S (16.32)  \ncould be compensated by a backward reaction  \nOCS + S → CO + S 2 (16.33)  \nthat becomes more essential with altitude and may account for net OCS loss at upper 20x km, as discussed below. In contrast, the modeling of Bierson and Zhang (2020) indicates OCS production by reaction (16.31) at 10 -25 km and by reaction (16.32) below 10 km. The importance of reaction (16.32) was discussed by von Zahn et al. (1983) and Prinn (1985a, b). Then, its feasibility was demonstrated experimentally at Venus' near-surface temperatures by Hong and Fegley (1997a). Fegley (2014) suggests reaction (16.32) as the most likely pathway of OCS formation. The reaction between S2 and CO2 is less efficient in OCS production (Krasnopolsky 2013). Formation of OCS through the reduction of SO2 by CO (Eq. 16.24) may not be significant because of its low rate assessed experimentally (Ferguson 1918; Hong and Fegley 1997a) and through kinetic evaluations (Fegley et al.  1997b). In addition to gas-phase reactions, OCS could be supplied through interactions of CO2 with sulfide minerals (Lewis 1970; von Zahn et al. 1983; Fegley and Treiman 1992a; Fegley et al. 1995) ( Sect. 16.3.2.3 ).  \nThe formation of OCS could be enhanced by the catalytic effect of dust in the supposed atmospheric haze at altitudes of 1 -2 km (Grieger et al. 2004) and by minerals in a permeable surface layer. Surface catalysis could promote chemical equilibration between OCS, CO, SO2, and CO2. To be consistent with  \n<!-- image -->  \nobserved concentrations of gases (Table 16.1), the equilibration is only feasible at the conditions of modal planetary radius (Krasnopolsky and Pollack 1994; Zolotov 1996). Calculations of gas-phase chemical equilibria for the surface atmosphere partially represented in Table 16.2 suggest x OCS within 5 -30 ppmv as follows, 20 ppmv (Krasnopolsky and Parshev 1981), 3 ppmv (Zolotov 1985), 13 ppmv (Krasnopolsky and Pollack 1994), 28 ppmv (Zolotov 1996), 16 -29 ppmv (Fegley et al. 1997b), 5 ppmv (Hong and Fegley 1997a), 20 ppmv (Krasnopolsky 2013), and 9 ppmv (Jacobson et al. 2017a).  The discrepancy can be attributed to variations in the chosen chemical systems, reactions, utilized x CO, x SO2, and x S2, temperature, and pressure among different studies. Lower x CO of ~8 ppmv obtained through extrapolation of Venera 12 data (Table 16.1, Gel'man et al. 1980) to the surface (Fegley et al. 1997a, b) corresponds to 4 -5 ppmv OCS at equilibrium (Hong and Fegley 1997a; Model 9 in Table 16.2). Equilibrium calculations using measured S3 (Zolotov 1985; Model 10 in Table 16.2) suggest only a few  \nppmv OCS.", 'Consumption of OCS': 'Photochemical models do not indicate net production of OCS in the mesosphere and clouds because of photolysis and overall oxidizing conditions caused by photolysis. According to Krasnopolsky (2013), net OCS consumption in the middle atmosphere occurs with net reaction  \nOCS + 2CO2 → SO 2 + 3CO, (16.34)  \nwhile net decay to CO and Sn is estimated to be insignificant. Actual pathways include photodissociation and interaction with O produced through photolysis of CO2, SO2, H2O, SO, and SO3  \nOCS + hv → CO + S (16.35)  \nOCS + O → SO + CO (16.36)  \n(Krasnopolsky 2012; Zhang et al. 2012). Zhang et al. (2012) and Bierson and Zhang (2022) also considered OCS reactions with Cl, ClS, S, and NO3 as possible loss routes and inferred reaction (16.35) as the main pathway of CO production at 65 -75 km.  \nKrasnopolsky (2007) estimated 14 ppbv OCS at 58 km, which reflects a primary OCS consumption within clouds that follows eddy diffusion from the lower atmosphere. The modeled fate of OCS in upper clouds by Krasnopolsky (2012) agrees with the cloud top mixing ratio of ~3 ppbv at 65 km and with a prominent  \n64  \n<!-- image -->  \ndecrease in x OCS with altitude at 64 -72 km observed by Krasnopolsky (2008, 2010b). A significant consumption of OCS at and above the cloud top is modeled by Zhang et al. (2012), Zhang and Bierson (2020), Pinto et al. (2021), and Dai et al. (2024).  \nMost models do not expect an increase in x OCS with altitude above the cloud top (Mills 1998; Krasnopolsky 2012; Zhang et al. 2012; Bierson and Zhang 2020; Dai et al. 2024). The detection of a robust positive gradient at 65 -100 km and ~1 ppmv OCS at 100 km (Mahieux et al. 2023, Table 16.1) is inconsistent with modeling results for mesospheric OCS. Mahieux et al. (2023) noted that the inconsistency may indicate a source for OCS in the upper mesosphere.  Here, I propose a supply of FeS and S-bearing organic matter with cometary dust (Carrillo-Sánchez et al. 2020) as a source of abundant mesospheric OCS, CS, and CS2. This explanation aligns with the observed OCS gradient above cloud tops, the sporadic detections of OCS at low latitudes, and its relatively stable detection pattern in higher latitudes (Mahieux et al. 2023). Considering the short photolysis lifetime of mesospheric OCS (~6 hours), the observations of Mahieux et al. (2023) imply a steady supply of space materials. However, the observed variability in x OCS could indicate fluctuations in the supply of these materials.  \nAtmospheric models of Krasnopolsky and Pollack (1994), Krasnopolsky (2007, 2013), and Bierson and Zhang (2020) suggest OCS supply through eddy diffusion from the near-surface atmosphere (0 -20 km) and its net consumption above ~25 km. This scheme agrees with strong OCS gradients inferred for lower 30s km ( Sect. 16.2.1.3 , Table 16.1). A significant OCS consumption is modeled through oxidation by SO3 at 30 -40 km. There is no consensus regarding the specific reaction pathways involved in this process (Eqs. (16.18) -(16.21)), and experimental data on reaction rates and mechanisms are needed to assess the thermochemical oxidation of OCS ( Sect. 16.3.1.2, 16.3.1.4 ).  \nPrinn (1975, 1978) first proposed OCS interaction with atomic sulfur (Eq. (16.33)) as a significant pathway of OCS consumption and S2 production in the middle and lower atmosphere. However, neither observation nor model suggests an elevated abundance of atomic S needed to consume OCS at 30 -36 km, where x OCS drops sharply with altitude. The modeling of Krasnopolsky and Pollack (1994) and Krasnopolsky (2007) does not indicate that reaction (16.33) is efficient at a low x Sn at 30 -36 km. To avoid the problem with Sn deficiency, Yung et al. (2009) suggested OCS interaction (16.33) with supposedly abundant atomic S produced by a soft-UV photolysis of S3 and S4 in the sub-cloud atmosphere  \n<!-- image -->  \nS4 + hv → S3 + S (16.37)  \nS3 + hv → S2 + S. (16.38)  \nWhether this mechanism works could be tested by assessing the abundance of Sn gases and by gaining better knowledge about UV irradiance below clouds. Krasnopolsky (2013) included photochemical reactions proposed by Yung et al. (2009) in his 2007 model and confirmed that oxidation of OCS by SO3 permits the principal OCS loss at 30 -36 km, consistent with the anticorrelation of OCS and CO and their opposite vertical gradients ( Sect. 16.2.1.3 ). However, models of Krasnopolsky (2013) and Bierson and Zhang (2020) show that reaction (16.33) significantly contributes to OCS loss below 30 km. Krasnopolsky (2013) demonstrated that reaction (16.33) reaches the highest efficiency at 28 km and causes OCS loss at upper 20s km, inferred through observations by Pollack et al. (1993). Results of Bierson and Zhang (2020) suggest that reaction (16.33) is accountable for OCS loss at ~5 -30 km. Reaction  \nOCS + M → CO + S + M (16.39)  \nis modeled as accountable for OCS loss below 5 km (Krasnopolsky 2007; Bierson and Zhang 2022), while the reaction  \nOCS + S2 → CO + S3 (16.40)  \nprovides a lesser contribution (Krasnopolsky 2007, 2013). OCS oxidation to CO on the surface of hematite, likely a secondary mineral ( Sect. 16.3.2.3 ), is too slow to influence OCS consumption, as Yung et al. (2009) estimate. Despite the assessed OCS loss below ~25 km by reactions (16.33), (16.39), and (16.40), the net OCS production is commonly modeled there, as discussed above.', '16.3.1.7  Sulfur Gases and Condensates': "Although sulfur gases (Sn) and condensates (Sx, S8) could play significant roles in atmospheric chemistry, observational data are limited ( Sects. 16.2.1.3, 16.2.1.4 ), and there is no consensus about their occurrence, abundance, speciation, and chemical pathways. Because of limited observational data on sulfur species, the current understanding of their behavior is based on models. The ambiguity of modeling results reflects the lack of rate constants of critical reactions.  \n<!-- image -->  \nSn gases could be produced above clouds (e.g., Eq. (16.35)) and through sub-cloud thermochemical reactions (e.g., Eq. 16.18), affected by photolysis by soft UV photons below 30 km (Eqs. (16.37), (16.38)), as well as consumed and formed in thermochemical reactions below 30 km (Eqs. (16.31), (16.32), (16.33), (16.39), and (16.40), Table 16.7). S2 could be produced in gas-mineral reactions ( Sect. 16.3.4.2 ) and be present in volcanic gases ( Sect. 16.4.3 ). Chemical equilibrium models for the sulfur system (San'ko 1980; Zolotov 1985; Krasnopolsky 2013) suggest increasing stability and relative abundances of heavier Sn gases with altitude with a dominance of S8 at high altitudes. Barsukov et al. (1980a, 1982a) and Dorofeeva et al. (1981) considered the phase diagram of sulfur concerning condensation. They noted the possibility of condensation at the cloud base at a total Sn abundance of ~0.2 ppmv suggested by gas-phase chemical equilibria at the surface (e.g., Models 4, 5 of Table 16.2). The phase diagram does not imply condensation at the cloud base at total Sn of ~0.01 ppmv inferred from Venera spectrophotometer data below 25 km (Moroz et al. 1981; San'ko 1980; Sect. 16.2.1.3 , Model 10 of Table 16.2) The bulk Sn content within and around clouds remains unknown, and it is unclear if sulfur condenses there. The major constraint is the diversity in model-based assessments on bulk content and sources of sulfur in clouds. Both photochemical (from above) and thermochemical (from below) sources are considered in the literature. There is more consistency in understanding Sn gases below 30 km, where S3 and S4 are constrained from observations ( Sect. 16.2.1.3 , Table 16.1), and the dominance of S2 in the near-surface atmosphere is suggested from chemical equilibrium assessments (Table 16.2). Bulk Sn content and x S2 of 0.1-0.4 ppmv correspond to chemical equilibria calculated from measured CO2, CO, and SO2 (Models 1-8 of Table 16.2). This bulk Sn content allows Sx condensation in clouds at ~50 km (Krasnopolsky 2013). The consideration of extrapolated x CO or observational x S3 data results in 0.010.02 ppmv S2 and bulk Sn in the near-surface atmosphere (Models 9 and 10) that could impede condensation in clouds.", 'Pathways of Gaseous and Condensed Sulfur in the Middle Atmosphere': "The fate of sulfur species within and around clouds has been considered by Prinn (1975, 1985a, b), Dorofeeva et al. (1981), Barsukov et al. (1982a), Young (1983), von Zahn et al. (1983), Krasnopolsky and Pollack (1994), Mills and Allen (2007), Zhang et al. (2012), Krasnopolsky (2012, 2016), Pinto et al. (2021), and Francés-Monerris et al. (2022). Figures 16.19 and 16.20 show that Sn could form and be consumed through multiple reactions. So far, prevailing chemical pathways and amounts of Sn produced in the middle atmosphere are insufficiently constrained. One reason for the ambiguity is the lack of data on  \n67  \n<!-- image -->  \nreduced S-bearing gases within clouds. Another reason is insufficient information on the kinetics of numerous Sn species reactions.  \nBefore the Pioneer Venus mission, Prinn (1975, 1978) considered the formation of S and S2 through the photolysis of OCS via (16.35) and its reaction (16.33) with S and via the interaction of two HS molecules formed through photolysis of H2S. His results show that the presence of oxidants such as O2 in upper clouds and above limits the net production of Sn from supposedly abundant OCS and H2S. In his models, atomic S is consumed by reaction (16.28) with O2 at and above the cloud top and via (16.33). This produces S2 at deeper O2-poor conditions and could lead to Sx upon polymerization and condensation. von Zahn et al. (1983) and Prinn (1985a, b) further emphasized the net formation of Sn species through OCS photolysis and HS + HS interactions in upper clouds. Mills et al. (2007) mentioned the possible formation of free sulfur via disproportionation of SO2 to Sn and SO3 at cloud tops.  \nMills and Allen (2007) considered S2 formation in upper clouds through reaction cycles involving Clbearing species, SCl, SCl2, S2Cl, ClCO, Cl, and Cl2. Their models suggest the formation of atomic S mainly via photolysis of SO and SCl, followed by S2 formation through the dissociation of SCl2 and SCl. The consumption of S and S2 is modeled through photochemical reactions in upper clouds with the participation of SCl, S2Cl, SCl2, and Cl-O species. At cloud top (63 -70 km), S and S2 gases are supposedly consumed by reactions (16.28) and (16.41), and reactions (16.42) and (16.43) are modeled as responsible for losses in more reduced upper clouds at 58 -63 km  \nS2 + O → SO + S (16.41)  \nS + Cl2 → SCl + Cl (16.42)  \n2S2 + M → S 4 + M. (16.43)  \nMills and Allen (2007) noted a significant negative effect of x O2 and the O2/Cl ratio on net Sn production and modeled increasing losses of S and S2 with altitude. Building on the work of Prinn (1975), Mills et al. (2007) posited that the fate of sulfur -whether it transforms into sulfuric acid or Sx condensate -is governed by a competitive process: the oxidation of atomic S by O2 via (16.28) versus its reaction (16.33) with OCS.  \n<!-- image -->  \nThe production and fate of free sulfur in the middle atmosphere have been considered in Krasnopolsky's (2012) and Zhang et al. (2012) models. Their nominal models for clouds at 47 -60 km suggest mixing ratios of Sn (mainly S8) of 10 -12 to 10 -9 that confine condensation. Zhang et al. (2012) demonstrated that the photolysis of OCS (16.35) and the interaction between OCS and S (16.33) could significantly contribute to the production of S and S2 at 58 -62 km if there is a substantial concentration of OCS. This OCS abundance could be attributed to increased eddy diffusion from the lower atmosphere or other factors that are not yet justified. Krasnopolsky (2012) showed that formation and photolysis of (SO)2 (Eqs. (16.25), (16.26)) are more effective in S formation than pathways listed above, though (SO)2 is not assessed as abundant in this and other photochemical models (Krasnopolsky 2018; Bierson and Zhang 2020; Pinto et al. 2021). In Krasnopolsky's (2012) model, S is readily lost by reaction (16.28) with O2, which restricts the accumulation of Sn gases sufficient to form Sx aerosol in upper clouds. The formation of atomic S in clouds via thermochemical reduction of SO2 by CO by (16.12) is considered negligible at upper clouds' overall oxidizing conditions (Krasnopolsky 2013). In these scenarios, condensed sulfur is absent from upper clouds, zero net transport of free sulfur to the sub-cloud region is expected, and possible condensed Sx in clouds is sourced from beneath.  \nAssuming a sufficient formation of atomic S in initial reactions (e.g., (16.26)), several research groups addressed polymerization pathways to constrain condensation within clouds. Mills and Allen (2007) modeled the formation of heavier Sn species via reactions (16.43) and (16.44),  \nS + S2 + M → S 3 + M (16.44)  \nIn contrast to the results of Zhang et al. (2012) and Krasnopolsky (2012, 2018) that do not predict much (SO)2 and Sn in upper clouds, photochemical models of Pinto et al. (2021) and Francés-Monerris et al. (2022) considered reactions of (SO)2 that could produce S3, S4, and polysulfur species (Fig. 16.19). Their results for C-O-N-Cl-S species at 55 -105 km suggest S2 as the most abundant Sn gas at the cloud tops (Fig. 16.21). Pinto et al. (2021) modeled the photodissociation of (SO)2 by reactions such as (16.26) and  \ncis-(SO)2 + hv → SO + SO → S 2 + O2 (16.45)  \nand polymerization of S2 to Sn as a source of condensed sulfur. Quantum-chemistry computations and photochemical models of Francés-Monerris et al. (2022) demonstrated that S2 may not be a primary product of (SO)2 photolysis ((16.26), (16.45)) and modeled S2 formation via reactions.  \n<!-- image -->  \n(SO)2 + SO → S 2O + SO2 (16.46)  \nClS + SO → S 2O + Cl (16.47)  \nS2 O + SO → S 2 + SO2 (16.48)  \nthat could allow the formation of Sx species from SO and cis-(SO)2 through the S2O intermediate and maintain a significant fraction of condensed Sx in cloud aerosols.  \nModels for the middle and lower atmosphere (Krasnopolsky 2007, 2013; Bierson and Zhang 2020) suggest a variable bulk sulfur content with altitude due to competitive consumption and production reactions.  For the cloud base at 47 km, Krasnopolsky (2007) modeled major Sn gases as follows: S8 (80%), S7 (8.5%), and S6 (10.5%).  He estimated Sn density close to saturation at cloud tops and complete evaporation of sulfur aerosol at the cloud deck of ~48 km. The model of Krasnopolsky (2013) reveals the formation of Sx aerosol through condensation in lower clouds (at ~50 km). That model excludes Sx from being an effective UV absorber(s) observed in upper clouds ( Sect. 16.3.1.8 ) if sufficient Sx does not form there.", 'Sulfur Species in the Lower Atmosphere': "If condensed sulfur is present in clouds ( Sect. 16.2.1.4 ), the subsidence of aerosol particles by gravity and circulation removes native sulfur from the middle atmosphere. Evaporation of Sx within a few km below the cloud deck would be followed by sequential thermal decomposition of S8 to lower-mass allotropes, and S2 dominates at lower altitudes (Bierson and Zhang 2020; Krasnopolsky 2013; San'ko 1980; Zolotov 1985, Figs. 16.15, 16.17) and the surface (Table 16.2).  \nModels suggest the thermochemical formation of atomic S below 30 km by photolysis of S3 and S4 (Eqs. (16.37), (16.38)) (Krasnopolsky 1987, 2013), the interaction of SO molecules (Eq. (16.23)) formed through SO2 + CO interaction (16.22) below 16 km that together corresponds to net S production by (16.12) (Krasnopolsky 2013), S2 interaction (16.32) with CO below 10 km (Bierson and Zhang 2020), and OCS dissociation by (16.39) below 5 km (Krasnopolsky 2007, 2013; Bierson and Zhang 2022).  S2 is modeled as a byproduct of OCS oxidation by SO3 at min-30s km (Eqs. (16.18) -(16.20)) (Krasnopolsky and Pollack 1994; Krasnopolsky 2007, 2013), though reaction (16.20) might be inefficient because of the short thermochemical life of (SO)2 (Bierson and Zhang 2020). S2 forms through OCS oxidation by S by  \n<!-- image -->  \n(16.33) at 5 -30 km (Krasnopolsky 2013; Bierson and Zhang 2020), via S3 photolysis via (16.38) at 3 -29 km, and through thermal decomposition of S4 assessed by Krasnopolsky (2013). S3 forms through S4 photolysis (16.38) at ~18 km and via S2 -OCS interaction (16.40) below 5 km (Krasnopolsky 2013).  \nThe modeled and discussed consumption of Sn gases included thermochemical oxidation by SO3 (Mills and Allen 2007), photochemical dissociation by soft UV photons (Eqs. (16.37), (16.38)) (Yung et al. 2009; Krasnopolsky 2013, Bierson and Zhang 2020), and reactions (16.31) and (16.32) with CO. The loss of Sn gases through the reduction of SO3 to SO2 is not supported by Krasnopolsky's (1987, 2013) and Bierson and Zhang's (2020) models. A photolysis of S3 and S4 to produce abundant Sn in the sub-cloud atmosphere proposed by Yung et al. (2009) has no observational support. Krasnopolsky (2013) and Bierson and Zhang (2020) included their models' kinetics of S4 photolysis by 16.37. They did not infer abundant Sn that strongly affects OCS loss in the mid-30s km ( Sect. 16.3.1.6 ). Models on Krasnopolsky (2007, 2013) and Bierson and Zhang (2020) suggest consumption of atomic S through oxidation of OCS by (16.33) below 30 km and via reduction of CO to OCS by (16.31) below 25 km. A lot more abundant S2 is partially consumed through (16.32) with CO in the near-surface atmosphere. Reaction (16.33) could compensate for and overwhelm S2 loss at higher altitudes. Net loss of Sn gases from the lower atmosphere occurs through transport to the middle atmosphere. In Krasnopolsky's (2013) model, an upward eddy diffusion of Sn is followed by condensation in lower clouds at 50 km.", '16.3.1.8  Unknown Blue-UV Absorber in Clouds': "In addition to gaseous SO2 , Venus' spectra suggest an unknown absorber or absorbers in the UV and blue (320 -500 nm) spectral ranges (Esposito 1980). The featureless spectra of Venus in the near-UV range and in situ nephelometry data from Pioneer Venus and Venera 14 probes suggest a condensed absorber (Ekonomov et al. 1984) occurring in upper clouds above 57 -60 km ( Sect. 16.2.1.4 ). Among various potential compounds, such as FeCl3, NO, and Cl2, both gaseous and condensed S-bearing species have been posited to explain the absorption at short wavelengths. The spatial correlation of SO2 and an unidentified UV-blue absorber, deduced from Akatsuki UV observations, indirectly indicates the involvement of an S-bearing absorber in atmospheric processes alongside SO2 (Yamazaki et al. 2018). The considered candidate S-bearing species are as follows, elemental sulfur (Sx) (Hapke and Nelson 1975; Young 1983; Krasnopolsky 2016), S3 and S4 (Toon et al. 1982), S2O (Hapke and Graham 1989; Na  \nand Esposito 1997), CS2 (Barker 1978; Young 1978), SCl2 (Krasnopolsky 1986), NOHSO4 (Watson et al.  \n<!-- image -->  \n1979), irradiated S2O (Pérez-Hoyos et al. 2018), (SO)2 gas (Frandsen et al. 2016; Pérez-Hoyos et al. 2018), and ferric sulfates (Jiang et al. 2024). Many of these species are insufficiently consistent with the whole UV-blue spectra of Venus, observed cloud properties, and atmospheric and chemical models (Titov et al. 2018; Pérez-Hoyos et al. 2018).  \nCondensed native sulfur is stable in sulfuric acid clouds (Young 1983). It could be present in lower clouds based on Vega in situ data ( Sect. 16.2.1.4 ). However, the models of Zhang et al. (2012) and Krasnopolsky (2012) do not suggest the formation and accumulation of abundant Sn species in upper clouds ( Sect. 16.3.1.7) . Krasnopolsky (2016) calculated the condensed Sx profile through clouds and concluded that sulfur cannot be a significant UV absorber. Spectra of S4, SCl2, NOHSO4, and S2O do not match the whole UV-blue spectrum of Venus, especially between 0.4 µm and 0.5 µm (Pérez-Hoyos et al. 2018), though a mixture of absorbers may be consistent with observations. The presence of S4 in such a mixture would explain absorption at the longer wavelengths (Pérez-Hoyos et al. 2018), though S4 is inferred from the UV absorption pattern only below ~20 km (Krasnopolsky 2013). Another concern with Sx serving as the UV absorber is its low solubility in sulfuric acid, suggesting its accumulation on the surface of aerosol droplets (Young 1983) and separate aerosol particles. Consequently, this could diminish the glory optical phenomenon of upper clouds, as shown by Petrova et al. (2018).  \nAlthough three SO dimers ( Sect. 16.3.1.3 ) are strong absorbers in the blue and UV wavelengths (Frandsen et al. 2016; Pérez-Hoyos et al. 2018), photochemical models of Zhang et al. (2012), Krasnopolsky (2012, 2018), and Zhang and Bierson (2020) do not predict sufficiently abundant (SO)2 to be the main UV absorber in upper clouds, consistent with lack of (SO)2 detection. Nevertheless, upper cloud models of Pinto et al. (2021) and Frances-Monerris et al. (2022) suggest reactions of (SO)2 (Eqs. 16.45 -16.48) that could produce abundant S2O, polysulfur oxides, S3, S4, and condensed Sx species (Fig. 16.19) that all absorb in the blue-UV range.  \nJiang et al. (2024) demonstrated that the interaction of iron with concentrated sulfuric acid leads to the formation of rhomboclase, (H5O2)Fe(SO4)2·3H2O, and acid ferric sulfate (H3O)Fe(SO4)2. The authors showed that a mixture of these solid phases with liquid Fe 3+ -bearing sulfuric acid sufficiently matches the spectrum of Venus. However, the presence of iron in clouds suggested from in situ measurements (Petryanov et al. 1981b; Andreichikov et al. 1987; Zolotov et al. 2023; Mogul et al. 2024, Sect. 16.2.1.4 ) needs confirmation, and it is unclear if the supply of iron from space (Fe-metal, FeS, Carrillo-Sánchez et  \n<!-- image -->  \nal. 2020) and/or the surface compensates for the sinking of ferric sulfates toward the surface where they irreversibly decompose to stable hematite, α -Fe2O3 (Zolotov 2021; Zolotov et al. 2023). Today, there is no consensus on the composition of the UV-blue absorber in clouds.", '16.3.1.9  Polysulfur Oxides, Hydrogen Sulfide, and Carbon Sulfides': 'None of the reduced S-O and S-H gases are detected in the middle and lower atmosphere, and CS and CS2 are only observed at and above cloud tops (Table 16.1). Evaluation of the fate of these gases is strongly model-dependent, and there is no consensus regarding limited modeling results.', 'Polysulfur Oxides (SxO)': "Although no observational data on SxO gases are available, S2O is a potential condensable species in the middle atmosphere that absorbs in the UV range ( Sect. 16.3.1.8 ). Yung and DeMore (1982) first discussed the formation of S2O in the middle and upper atmosphere via (SO)2 + SO interaction (16.46) and consumption by reactions  \nS2O + hv → S + SO  \n(16.49)  \nS2 O + S → S 2 + SO (16.50)  \nS2 O + O → 2SO (16.51)  \nat 58 -100 km. Na and Exposito (1997) modeled the formation of an abundant S2O in upper clouds using reaction kinetics of (16.46), (16.49), and (16.51). The accumulation of S2O was attributed to its gradual depletion via (16.51). However, the models of Mills (1998) imply several orders of magnitude smaller xS2O than the evaluations of Na and Esposito (1997) because of the faster rate used for (SO)2 loss by back reaction of (16.25). Mills's (1998) results do not suggest that S2O is a major UV absorber in upper clouds.  \nThe photochemical modeling of Pinto et al. (2021) and Frances-Monneris et al. (2022) suggest reactions of SO and SO dimers as an essential source of S2O and other SxO species in the middle and upper atmosphere. Pinto et al. (2021) modeled the formation of S2O through the interaction of SO with (SO)2 isomers (e.g., by (16.46)) and ClS via (16.47) and by reaction  \n<!-- image -->  \nClO + S2 → S 2O + Cl (16.52)  \n(Fig. 16.19). In their models, x S2O reaches a ppbv level at the cloud top (~65 km) and becomes less abundant in upper clouds (~58 km) and the mesosphere. Calculated profiles of S3O and S4O show the same pattern at lower mixing ratios. At 100 km, the estimated mixing ratios of all SxO species are below 10 -15 . Frances-Monneris et al. (2022) modeled the formation of S2O through (16.46) and (16.47) and the consumption of S2O through interaction with SO by (16.48) at 58 -110 km. In their models, S2O reaches a maximal concentration at ~63 km (Fig. 16.21). At 60 km, the modeled densities of S2O, cis-(SO)2, and S2 formed by (16.48) are equal. The results suggest an essential S2O role as an intermediate species in forming Sn and other possible UV absorbers in upper clouds ( Sects. 16.3.1.7, 16.3.1.8 ).", 'Hydrogen Sulfide': 'Although H2S has not been firmly detected ( Sect. 16.2.1.3 ), it has been discussed as a key reduced gas by Prinn (1978, 1985a, b) and von Zahn et al. (1983). They suggested H2S formation through interaction between surface pyrite and water vapor ( Sect. 16.3.2.3 ) and via thermochemical reactions in the deep atmosphere and consumption by photochemical dissociation and oxidation in the middle atmosphere by net process  \nH2S + CO2 → H 2O + 2CO + SO2. (16.53)  \nYung and DeMore (1982) discussed H2S production and loss pathways via photolysis and interaction with atomic H. Still, they did not include corresponding reactions in the model because of insufficient information on reaction rates. The models of Krasnopolsky (2007, 2013) incorporate rates of thermochemical reactions involving H2S and estimated x H2S in the lower and middle atmosphere. It is assumed that H2S equilibrates with other gases (H2O, SO2, CO, CO2) at the surface. Modeling results suggest the thermochemical formation of HS radicals by reactions  \nH + OCS → CO + SH (16.54)  \nH2 S + S → SH + SH. (16.55)  \nH2S forms through reactions of HS with HCl, H2O, HSCl, and H2 and by back reaction (16.55). The relatively fast reaction  \n<!-- image -->  \nHS + HCl → H 2S + Cl  \n(16.56)  \nprovides a major contribution. The formation of H2S is balanced by (16.55) and by the flux toward clouds where H2S is oxidized. Maximal yields of reactions (16.55) and (16.56) are modeled at 18 km and 26 km, respectively. The modeled decrease in x H2S from assumed 150 ppbv at 0 km altitude to modeled 32 ppbv at the cloud deck at 47 km reflects the upward flux. The models of Bierson and Zhang (2020) and Dai et al. (2024) result in a constant x H2S in the lower and middle atmosphere and a major drop above ~65 km. Their results are consistent with the upper limit for x H2S at the cloud top reported by Krasnopolsky (2008) (Table 16.1). Neither Pioneer Venus Large Probe (Hoffman et al. 1980) nor Venera 13/14 entry probe (Mukhin et al. 1983) data on H2S are supported in these models. Most chemical equilibrium models for the near-surface atmosphere that assume 20 -30 ppmv H2O suggest x H2S of 0.01 -0.15 ppmv (Table 16.2).', 'Carbon Sulfides': 'Although abundant CS and CS2 are detected at and above cloud top (Table 16.1), they have yet to be included in atmospheric models, and the observed abundances of these reduced species in the overall oxidizing mesospheric environments remain to be understood. The increasing mixing ratios of carbon sulfides with altitude correlate with SO3, SO2, SO, and OCS trends in the upper mesosphere ( Sect. 16.2.1 ). As noted with respect OCS ( Sect. 16.3.1.6 ), one possible explanation of mesospheric CS and CS2 is their formation through photochemical alteration of extremely C-rich cometary dust (see Table 16.12 for carbon flix) containing up to 50% organic matter by mass (Bardyn et al. 2017).', '16.3.1.10  Chlorine-bearing Sulfur Gases': "Formation and consumption of ClSO2, SO2Cl2, SOCl, ClSO4, SCl, S2Cl, SCl2, and HSCl have been considered since the early 1980s. Evaluations and models suggest the formation of these species in the middle atmosphere through HCl photolysis and consumption in the sub-cloud atmosphere back to HCl and Sbearing gases. Strattan et al. (1979) and Yung and DeMore (1982) first discussed the role of reaction  \nSO2 + Cl + M → Cl SO2 + M (16.57)  \nin the chemistry of Cl and ClOx species that may affect the oxidation of SO2 in upper clouds. DeMore et al. (1985) experimentally demonstrated anoxic formation of SO2Cl2 through photolysis of mixtures of SO2  \n<!-- image -->  \nand Cl2. Inferred reactions in the S-O-Cl system included the formation of intermediate ClSO2 via (16.57) and its conversion to SO2Cl2 by reaction  \nClSO2 + ClSO2 → SO2Cl2 + SO2. (16.58)  \nAtmospheric models with corresponding reactions suggested SO2Cl2 is a key Cl-bearing species in upper clouds with a mixing ratio of 4 ppmv. Oxidation of ClSO2 by O2 to ClSO4 was indicated as a critical pathway for O2 consumption and SO2 oxidation to H2SO4 ((16.11), Sect. 16.3.1.2 ). Although subsequent modeling by Krasnopolsky (2012) and Bierson and Zhang (2020) did not support a significant role of Clbearing species in SO2 oxidation, the role of ClSO2 remains unclear because reaction rates are uncertain.  \nMills (1998) considered reactions with SClOx and SClx within and above upper clouds and provided the first quantitative estimation of their chemistry. SCl2, SCl, S2Cl, and S2Cl2 (least abundant) are predicted to exist only within the upper cloud layer at ~56 -60 km. ClSO2 is expected to be more abundant than SxCly gases. The model shows that ClSO2 could serve as a buffer for SO2, which slows its oxidation rate to H2SO4. At the same time, SxCly species could be intermediaries in the polymerization of atomic S to Sn. Mills and Allen (2007) further discussed and evaluated the role of SCl, S2Cl, and SCl2 as intermediate species that affect Sn gases. In their model, the formation of S2 occurs via S2Cl and SCl ( Sect. 16.3.1.7 ). As in the work of Mills (1998), the formation of S-Cl gases was modeled to appear in the upper cloud layer that is depleted in O2.  \nZhang et al. (2012) modeled the fates of ClSO2, SOCl, SCl, ClS2, SCl2, and S2Cl2 in the middle and upper atmosphere at 58 -110 km. The formation of the most abundant ClSO2 is modeled by (16.57), in which M is CO2. ClSO2 is consumed via reactions with O, S, S2, H, Cl, SO, SCl, and ClSO2 (Eq. (16.58)), producing SO2. Estimated x ClSO2 and x SOCl almost reach a ppbv level at the cloud top of 65 km and are less abundant at higher and lower altitudes. At 65 km, SCl2, SCl, ClS2, and S2Cl2 concentrations are within 1 -100 pptv and drop at higher altitudes due to oxidation. SCl2 is more abundant than ClSO2 in O-depleted upper clouds at 58 -60 km.  \nKrasnopolsky (2012, 2013) modeled ClSO2, OSCl, SO2Cl2, SCl2, and SCl at 47 -112 km and 0 -47 km. SO2Cl2 is estimated to be HCl's most abundant photochemical product, with a mixing ratio of 8 ppbv at 47 km and 30 ppbv at 68 km. SO2Cl2 forms via (16.57) and (16.58). Its net production in upper clouds proceeds via photolysis of CO2 via (16.1) and HCl and by reaction  \n<!-- image -->  \n2HCl + SO2 + CO2 → SO2Cl2 + H2O + CO. (16.59)  \nThe modeling suggests a net supply of SO2Cl2 to the lower atmosphere and other photochemical products, CO and H2SO4(g). Below 30 km, SO2Cl2 is consumed by reactions  \nSO2Cl2 + S → ClSO 2 + SCl (16.60)  \nSO2Cl2 + Cl → ClSO 2 + Cl2  \n(16.61)  \nThe assessed x SO2Cl2 is 7 pptv at the surface, and a detectable constant concentration of 3.5 ppbv is estimated above ~30 km because of the lack of sinks and sources of the gas (Fig. 16.17). At 0 -25 km, the major Cl-S-bearing gas is ClSO2 at ~20 pptv, while HCl is the most abundant Cl-bearing species at 0.5 ppmv in the lower atmosphere.  \nBierson and Zhang (2020) modeled SO2Cl2 as the most abundant species among seven SxClyOz and SxCly gases at 30 -80 km with a mixing ratio of up to 0.1 -10 ppbv at 70 -80 km. x SO2Cl2 strongly decreases at lower altitudes in their model. Below 30 km, ClS2 is estimated to be the most abundant S-Cl-bearing gas with a mixing ratio of up to 1 ppbv at 0 km. Similarly, Dai et al. (2024) assessed SO2Cl2 as the major S-Clbearing gas at 65 -90 km at the pptv level. In the middle and lower atmosphere, they reported ClS2 as the most abundant gas at the ppbv level, followed by SCl2, S2Cl2, and HSCl existed at the pptv level. The understanding of the fate of Cl-S-bearing sulfur species is hindered by the absence of detection and the generally indeterminate rates of gas-phase interactions.", '16.3.2.1  Approaches, Constraints, and Major Pathways': 'Chemically active gases in the hot and dense near-surface air (Tables 16.1, 16.2) make gas-solid type chemical reactions unavoidable. CO2 and H2O are probable contributors as oxidizing agents of ferrous iron in exposed materials such as silicates, sulfides, oxides, and silicate glasses (e.g., Barsukov et al. 1982c; Fegley and Treiman 1992a, b; Fegley et al. 1997a; Zolotov 2018). However, the destiny of Sbearing gases is more intricate and less comprehended. The common vision is that S-bearing solids play a critical role in the atmosphere-surface system (Fig. 16.15; Von Zahn et al. 1983; Prinn 1985; Fegley and Treiman 1992a; Fegley et al. 1995) and in the physical-chemical evolution of that system since the global  \n<!-- image -->  \nvolcanic resurfacing event. It has been suggested that chemical trapping and release of gases at atmosphere-surface interfaces affect both the abundance and speciation of SO2, OCS, H2S, and Sn and trace non-S species such as CO throughout the atmosphere. The gas-solid interactions of S-bearing compounds could be important for maintaining sulfuric acid clouds and, therefore, for the thermal and chemical structure of the whole gas envelope (Fegley and Prinn 1989; Bullock and Grinspoon 2001). Permeable geological materials (minerals and glasses in bedrocks and products of their physical and chemical weathering) likely contain more sulfur mass than the atmosphere (Lewis and Kreimendahl 1980; Zolotov and Volkov 1992, Sect. 16.4.4.1 ) and may control the abundance and speciation of atmospheric sulfur via gas-solid reactions. Robert Mueller and John Lewis established the concept of chemical equilibria between chemically active atmospheric gases and crystal materials on Venus (Mueller 1963, 1964, 1965; Lewis 1968, 1970; Lewis and Kreimendahl 1980). Subsequent investigations supported the chemical equilibration of some trace atmospheric gases with crustal materials that could have been established through gas-solid interactions over time. This section considers gas-solid reactions and equilibria at the present temperature, pressure, and supposed gas composition (Tables 16.1, 16.2) below 10 km, the highest elevation on Venus above the 6052 km radius. Aspects of gas-solid interactions in the past and future are discussed in Sect. 16.4.4 .  \nThe elevated bulk sulfur abundance in the otherwise basaltic composition of Venera and Vega samples (Table 16.3) implies an exogenic (atmosphere-sourced and/or cosmic) source of sulfur in surface materials. The consumption of atmospheric sulfur via gas-solid weathering reactions is the standard explanation for these data. Sulfur dioxide likely plays a significant role in such reactions, as inferred via evaluations of mineral stability and targeted experimental studies. These efforts led to lists of stable and unstable minerals with respect to SO2 and an array of possible solid (sulfates, silica phases, silicates, etc.) and gaseous (S2, CO) products of these interactions. Less conclusive and sometimes conflicting evaluations have been obtained for the stability and fate of metal sulfides. Some sulfides were considered unstable, releasing S-bearing gases into the atmosphere. Stable sulfides could stay intact in rocks and form at the surface via gas-solid reactions or condensation.  \nCurrently, both chemical equilibrium and experimental evaluations of gas-solid interactions that impact S-bearing compounds are limited by various factors. The major constraint is the lack of instrumental data on almost all trace gases below 10 -20 km (Table 16.1), which limits the ability to understand and model near-surface gas and assess key fugacities ( f O2, f S2) that affect the stability of minerals. Existing  \n<!-- image -->  \nmodels for gas composition (Table 16.2) set boundary conditions for atmospheric models (e.g., Krasnopolsky 2007, 2013; Bierson and Zhang 2020) and assume gas-phase equilibration, necessitating verification. The apparent lack of gas equilibrium above the modal planetary radius, as inferred from a comparison of equilibrium and supposed measurement-based concentrations (Krasnopolsky and Parchev 1979; Krasnopolsky and Pollack 1994; Zolotov 1996), complicates the understanding of directions and fates of gas-solid reactions on highlands. Assessments of gas-solid reaction pathways via evaluations of mineral stability are restricted by uncertainties of thermodynamic properties of pure solids and solid solutions, by the typical formation of metastable reaction products, and by rates of solid-state diffusion of metals that could be time-limiting factors of reactions at 650 -750 K. Applications of gas-solid reaction laboratory experiments to Venus are disadvantaged by a limited capability to measure the gas composition and fugacity ( f O2, f S2, f SO2, etc.) in reaction vessels, the common use of steel vessels that affect the gas composition, by the short duration of the experiments (less than a few months) relative to common reaction rates, and by the necessity to trade experimental pressure, temperature, and run time to explore advanced alteration. A challenge to both equilibrium and experimental evaluations, the phase and chemical composition of bedrocks and their weathering products are poorly constrained beyond the apparent mafic lava composition of vast plains and volcanic centers ( Sect. 16.2.2 ).  Similarly, information on the weathering products of mafic materials at the Venera and Vega landing sites is confined to the increased sulfur content (Table 16.3) and the red slope in the near-IR spectral range (Golovin et al. 1983; Shkuratov et al. 1987). These observations imply a coating by a fine-grained ferric phase, possibly hematite (Pieters et al. 1986). Current knowledge of atmospheric and surface conditions, thermodynamic properties of solids, and experimentally observed early stages of gas-solid reactions allows for making only general inferences on the fate of S-bearing compounds at the current atmosphere-surface interface.  \nPublications on sulfur in the atmosphere-surface interactions on Venus have been reviewed by Volkov et al. (1986), Fegley and Treiman (1992a, b), Fegley et al. (1992, 1997a), Zolotov and Khodakovsky (1989), Zolotov and Volkov (1992), Wood (1997), Zolotov (2015, 2018, 2019), Gillmann et al. (2022), and Filiberto and McCanta (2024). Since the mid-1960s, evolving views on reaction pathways have reflected new data and models on atmospheric composition, advances in numerical and experimental modeling methods, and a growth of modeling efforts in the last decade. Somewhat divergent conclusions regarding gas-solid pathways over the past six decades have arisen due to variations in gas  \n<!-- image -->  \ncompositions, temperatures, and pressures employed in theoretical analyses and numerical and experimental models ( Sects. 16.3.2.2, 16.3.2.3 ). Although there is no consensus on details, major pathways include gas-solid reactions of rock-forming minerals and glasses with SO2, OCS, and S2 (Fig. 16.15).  \nSO2 reacts with basaltic glasses to form sulfates of Ca and Na, which are notable in alteration products of alkaline samples ( Sect. 16.3.2.2 ). S2 gas forms via sulfatization reactions at strongly anoxic surface conditions (Table 16.2) through disproportionation of sulfur  \n3S(IV) (in SO2 gas) → 2S(VI) (in metal sulfates) + S 0 (in S2 gas) (16.62)  \nThe disproportionation of sulfur in the sulfurization of soils by SO2 had been suggested from experiments with calcite (Tarradellas and Bonnetain 1973), and Burnham (1979) proposed it for the alteration of SO2 with Ca silicates in terrestrial volcanic environments. As an example, anhydrite forms via alteration of Ca-rich pyroxenes,  \nCaSiO3 (in pyroxene) + 1.5SO2 (g) → CaSO 4 (anhydrite) + 0.25S2(g) (16.63)  \nPlagioclases, especially Na-rich compositions, are more stable with respect to sulfatization based on evaluations of mineral stability and experiments. The interaction of OCS with ferrous iron in minerals (silicates, pyrrhotite, magnetite, etc.) and glasses could result in pyrite formation. Gas-solid reactions could lead to the equilibration of some gases with mineral assemblages formed in weathering reactions. In such a case, solids could buffer near-surface gas chemistry, and further alteration of geological materials could be impeded. The concentration of all S-bearing gases could be buffered by magnetitepyrite and/or magnetite-pyrite-hematite mineral assemblages ( Sect. 16.3.2.3 ).', '16.3.2.2   SO2 Interactions with Minerals and Glasses': 'Hot SO2 in the near-surface air is modeled to be a significant gas that alters surface materials, leading to a net sink of atmospheric sulfur to secondary minerals and the release of gaseous by-products (Fig. 16.15). The current understanding is that the formation of metal sulfates via overall reaction (16.62) contributes to the net oxidation of surface materials and the reduction of atmospheric gases by decreasing the SO2/S2 ratio. Considering S-rich compositions of surface compositions (Table 16.3), the  \n<!-- image -->  \nfate of SO2 at the atmosphere-surface interfaces has been constrained through evaluations of mineral stability concerning gas-solid type reactions and via experiments at simulated venusian and other conditions. Volkov et al. (1986), Fegley and Treiman (1992a, b), Fegley et al. (1992, 1997a), and Zolotov (2015, 2018, 2019) reviewed early works on SO2-mineral interactions, and this section devoted more attention to insufficiently addressed and recent publications and an array of new experiments.', 'Inferences from Mineral Stability': "Long before measurements of trace chemically active gases in the lower atmosphere, Mueller (1963, 1964) linked the stability of crustal minerals with atmospheric composition. He has demonstrated that gas-solid equilibria could affect atmospheric composition and that the measured composition of atmospheric gases could constrain the mineralogy of upper crustal materials that reacted with gas. Mueller (1965) estimated abundances of S-bearing atmospheric gases assuming their equilibration with crustal Ca- and Fe-bearing minerals in Venus' upper crust. He showed that chemical equilibration between wollastonite and anhydrite  \nCaSiO3 (wollastonite) + SO3(g) = CaSO4 (anhydrite) + SiO2 (quartz) (16.64)  \nat 700 K could only be achieved at p SO3 of 10 -13.8 bars. This implies the sulfatization of wollastonite at Venus' atmospheric p SO3, which exceeds the equilibrium value. However, assessing p SO2 at the CaCO3CaSO4 equilibrium required data on atmospheric f O2 to constrain the SO2/SO3 ratio. Mueller (1964, 1965) suggested that the oxidation state (e.g., f O2, CO2/CO, SO2/OCS, SO3/SO2 ratios) of a hot, coupled atmosphere-upper crust system could be controlled by chemical equilibria with Fe-bearing minerals and that the redox state of the system may not exceed that controlled by the magnetite-hematite (Fe3O4 -α -Fe2O3, Mt-Hem) mineral assemblage commonly used to maintain f O2 in petrological experiments (Eugster 1957)  \n2Fe3O4 (magnetite) + CO2(g) = 3Fe2O3 (hematite) + CO(g) (16.65)  \n4Fe3O4 + O2(g) = 6Fe2O3 (16.66)  \nMueller (1965) demonstrated that the magnetite-pyrite (Mt-Py) equilibrium  \n3FeS2 (pyrite) + 2O2(g) = Fe3O4 + 3S2(g) (16.67)  \n<!-- image -->  \nat f O2 controlled by (16.66) sets the S2/SO3 ratio that allows alteration of CaSiO3 to CaSO4 (Eq. (16.64)). Here, we note that p SO3 at (16.64) (10 -13.8 bars) at f O2 from (16.66) (10 -22.8 bars) corresponds to p SO2 and p S2 of 10 -4.8 and 10 -5.3 bars, respectively, assuming SO3-SO2-S2-O2 equilibria at 700 K. However, Mt-Py equilibrium (16.67) at p O2 controlled by the Mt-Hem equilibrium (16.66) corresponds to partial pressures of SO3, SO2, and S2 of 10 -11.2 , 10 -2.3 , and 10 -10.3 bars, respectively, that are in the stability field of CaSO4 at 700 K. Mueller (1965) stated that in a coupled atmosphere - upper crust system controlled by the Mt-Hem phase assemblage, sulfur would be stored in sulfates rather than in a CO2-rich atmosphere.  \nLewis (1968, 1970) considered gas-mineral equilibria of carbonates, silicates, and iron sulfides and oxides to link partial pressures of SO2, SO3, OCS, H2S, H2O, HCl, HF, CO, and CO2, surface mineralogy, temperature, and pressure to interpret remote sensing data on the upper Venus' atmosphere. The SO 2 mixing ratio evaluated for several temperature-pressure conditions did not exceed 0.3 ppmv (10 -4.44 bars SO2 at a total pressure of 120 bars) in the lower atmosphere. The latter p SO2 value corresponded to the CaCO3-CaSO4 equilibrium at atmospheric p CO2, supposedly controlled by the calcite-quartz-wollastonite equilibrium  \nCaCO3 (calcite) + SiO2 = CaSiO3 + CO2(g) (16.68)  \nthat was anticipated by Adamcik and Draper (1963) and Mueller (1963, 1964). Although the chemical consumption of atmospheric SO2 has yet to be discussed, the works of Muller (1965) and Lewis (1970) implied the depletion of any abundant (e.g., volcanic) atmospheric SO2 toward minute concentrations controlled by gas-mineral equilibria. Lewis and Kreimendahl (1980) noted that a mass of atmospheric sulfur corresponding to 100 ppmv of S-bearing gases (like current estimates, Table 16.1) could be consumed in a ~50 cm layer of surface dirt where all Ca-bearing phases are converted to anhydrite. This mass balance estimate supported Mueller's (1965) idea that hot upper crustal materials control atmospheric composition.  \nStability diagrams of S-, Ca-, and Fe-bearing minerals developed by Lewis and Kreimendahl (1980) further demonstrated that the stability of calcite concerning sulfatization required dominance of reduced S-bearing gases. This concurred with the estimations of Mueller (1965) and Lewis (1970) and with independent chemical models of the lower atmosphere developed by Prinn (1978). Lewis and Kreimendahl (1980) also considered a pyrite-calcite-anhydrite-wüstite f O2 buffer (10 -22.6 bars at 750 K)  \n<!-- image -->  \nconsistent with ~90 bars CO2 controlled by (16.68). Although Lewis and Kreimendahl (1980) suggested stable coexistence of CaCO3 and CaSO4 based on preliminary Pioneer Venus data, refined gas chromatography data on x SO2 (Oyama et al. 1980) implied instability of calcite.  \nSince 1980, sulfatization of Ca-bearing minerals in contact with 130 -185 ppmv SO2 reported by Oyama et al. (1980) and Gel'man et al. (1980) have been anticipated through an array of chemical equilibrium calculations of sulfatization reactions (e.g., (16.63)) and S-bearing multicomponent systems. Barsukov et al. (1980a, 1980b, 1982c) and Khodakovsky (1982) calculated equilibrium phase compositions of rockgas type systems in which rocks (basalts, rhyolite, etc.) were included in the models as their elemental composition (H, C, O, N, S, Cl, F, Na, Mg, Al, Si, K, Ca, P, Mn, Ti, and Fe). They used fixed fugacities of CO2, CO, SO2, and H2O (open system with respect to these gases), surface temperatures, and pressures obtained with the Pioneer Venus and Venera 12 probes. Equilibrium mineral assemblages calculated under Venus' lowlands commonly contained 5 -8 vol% anhydrite and occasional pyrite, and their sulfur content reached 11 wt%. These models showed that sulfatization of freshly exposed rocks and glasses could not affect plagioclase, alkali feldspars, quartz, enstatite, fluorapatite, rutile, and titanite that appeared in equilibrium assemblages. The presence of stable diopside in a CaSO₄ -free assemblage, calculated at minimal x SO₂ of 70 ppmv and maximal x CO of 42 ppmv CO from Gel'man et al. (1980), indicated its relatively high resistance to sulfatization (Barsukov et al. 1980a, 1982c). Barsukov et al.'s (1980b) calculations demonstrated greater stability for felsic igneous rocks (rhyolites, granites) than mafic rocks. Following Mueller's (1964, 1965) concept, Barsukov et al. suggested that their calculated solid phases represent the mineralogy of altered surface rocks and that the mineralogy of permeable surface materials buffers partial pressures of trace chemically active atmospheric gases (SO2, OCS, S2, CO, H2O, H2, etc.).  \nIn 1982, the anticipated sulfatization of some Ca-bearing minerals was supported by the detection of Srich solids at the landing sites of Venera 13 and 14 (Table 16.3). The S/Ca atomic ratios below unity indicated incomplete conversion of Ca-bearing phases to CaSO4. Barsukov et al. (1983) used these samples' elemental composition in chemical equilibria calculations in multicomponent gas-solid systems open with respect to SO2 and other atmospheric gases measured with PVLP and Venera 12 (Table 16.1). The resulting anhydrite-rich equilibrium assemblages contained more sulfur than measured in surface probes, suggesting incomplete (ongoing or hindered) sulfatization of surface materials. Equilibrium calculations without S-bearing atmospheric gases led to anhydrite-bearing assemblages consistent with  \n<!-- image -->  \nmeasured elemental compositions. Other secondary phases in the latter assemblages were enstatite, anorthite, oligoclase, minor diopside, nepheline (Venera 13), albite (Venera 14), magnetite, and quartz, as summarized by Volkov and Khodakovsky (1984) and Volkov et al. (1986).  \nZolotov (1985) noted that sulfatization on Venus could release S₂ gas (Eqs. 16.62 and 16.63). This inference was based on experiments involving the anoxic sulfatization of calcite (Tarradellas and Bonnetain 1973), which occurs via the formation of a sulfite intermediate  \nCaCO3 + SO2 (g) → CaSO 3 + CO2(g) (16.69)  \nCaSO3 + 0.5SO2 (g) → CaSO 4 + 0.25S2(g) (16.70)  \nPhase equilibrium lines for silicates, carbonates, and sulfates of Ca, Mg, Mn, Na, and K, plotted in coordinates of f SO2 and f S2 (Fig. 16.22, Volkov et al. 1986; Zolotov 1985, 2018) demonstrated the instability of Ca-bearing pyroxenes (wollastonite, diopside, augite) and carbonates (calcite, dolomite) concerning sulfatization at all elevations at x SO2 and x S2 suggested from atmospheric measurements. Conditions in deep lowland trenches (e.g., at Dyana Chasma) and/or shallow permeable subsurface appeared close to those at the anorthite-anhydrite equilibrium  \nCa2Si2Al2O8 (anorthite) + 1.5SO2(g) = CaSO4 + Al2SiO5 (andalusite) + SiO2 (quartz) + 0.25S2(g). (16.71)  \nf SO2f S2 diagrams developed for 670 K and 52.1 bars (e.g., Zolotov 2018) suggest more favorable sulfatization at highlands, where Ca-rich plagioclase is unstable regarding SO₂ and S₂. However, Mg-rich silicates (such as enstatite and forsterite) and magnesite (MgCO3) should not react with SO₂ at any elevation.  \nBarsukov et al. (1986b, c, d) modeled equilibrium mineral assemblages at Venera 13, 14, and Vega 2 landing sites in a 16-component gas-solid system with consideration of thermodynamic data on binary solid solutions of Fe-Mg olivine, Fe-Mg orthopyroxene, Fe-Mg-Ca clinopyroxene, plagioclase, alkali feldspar, and scapolite. Mineral assemblages calculated at the contact with a CO2-CO-H2 O Venus' atmosphere contained 1.5 -8 vol% CaSO4, whereas solids equilibrated with a CO2-CO-H2O-SO2-Cl-HF atmosphere had 13 -18 vol.% CaSO4 (Table 16.8). The latter values corresponded to 3.0 -4.3 wt% sulfur in altered compositions and indicated a maximum possible amount of trapped sulfur. Besides anhydrite, mineral assemblages equilibrated with the atmospheric SO2 contained K-dominated  \n<!-- image -->  \nmicrocline (Na/K = 0.13/0.87), clinoenstatite (Fe/(Fe + Mg) < 0.1), magnetite, andalusite, kalsilite (Venera 13  and Vega 2), marialite, tephroite, quartz, and rutile. Diopside, plagioclase (An20-30), and less abundant anhydrite in assemblages equilibrated with CO2-CO-H2O atmospheres (Table 16.8) suggested incomplete sulfatization of Ca-bearing minerals at all three landing sites. Formed and/or primary alkali feldspars and Mg-rich silicates were suggested among chemical weathering products at the landing sites and elsewhere. Different sulfur contents in solid samples (Table 16.3) were interpreted in terms of varying exposure of surface rocks, and shorter sulfatization time was suggested for S-poor and least physically degraded rocks at the Venera 14 site (c.f., Zolotov and Khodakovsky 1989; Zolotov and Volkov 1992). Although calculations of multicomponent gas-solid systems have shown a significant dependence of evaluated equilibrium phase assemblages on gas composition and the thermodynamic data used, the constant presence of anhydrite ensures the sulfatization of surface materials.  \nUsing the same Gibbs free energy minimization code as in Barsukov's works, Klose and Zolotov (1992) reported the formation of Na-, K- and Ca-sulfates in equilibrium assemblages formed through the alteration of evolved (alkali- and silica-rich) igneous rocks. As in Barsukov et al. (1980b, 1982c), abundant alkaline feldspars and quartz in equilibrium phases indicated elevated stability of evolved silicate rocks concerning sulfatization. Independent calculations of chemical equilibria in gas-basalt multisystem conditions with a consideration of mineral solid solutions by Klose et al. (1992) indicated the formation of anhydrite in all calculated assemblages and stability of Na-rich plagioclase and Mg-rich silicates. Semprich et al. (2020) used the Perple\\_X code (Connolly 2005) to calculate comprehensive phase equilibria, including solid solutions, in the f S2f O2 fields in which rocks and gases were presented by Si-O-Ti-Al-Fe-Mg-Ca-Na-K and C-O-H-S gas, respectively (Fig. 16.23). They considered diverse compositions of initial rocks and physical-chemical conditions at lowlands and highlands and calculated stability fields of silicates coexisted with anhydrite, pyrite, pyrrhotite, and Fe oxides in a range of f O2 and f S2 suggested for the near-surface atmosphere. Although Semprich et al. (2020) did not present results for silicate solid solutions, used a pyrrhotite composition of Fe0.88S, and only one sulfate (CaSO4), their calculations were roughly consistent with preceding models of multicomponent systems and considerations of individual gas-solid equilibria.  \nFegley and Prinn (1989) have demonstrated the instability of calcite through a consideration of equilibrium  \nCaCO3 + SO2(g) = CaSO4 + CO(g) (16.72)  \nat x SO2 and x CO obtained with the Pioneer Venus and Venera 12 probes (Table 16.1). They noted that equilibrium p SO2 in (16.72) is two orders of magnitude less than measured in the sub-cloud atmosphere. Subsequent comprehensive considerations of mineral stability by Fegley and Treiman (1992a, b) and Fegley et al. (1992) confirmed the instability of Ca-rich minerals (calcite, dolomite, diopside) and the stability of Mg-rich silicates in contact with supposed near-surface abundances of SO2 and CO. The discrepancy between a high and low f SO2 in the atmosphere and at mineral-anhydrite equilibria (e.g., Eq. (16.72)) was interpreted in terms of ongoing sulfatization.  \nZolotov (2018) revised evaluations of phase diagrams of mineral stability in coordinates of f SO2, f S2, and f CO and showed that plagioclase with an intermediate composition and activity of anorthite of 0.2 -0.6 could be in equilibrium with SO2, S2, and CO at the lowlands. Although the evaluation of anorthite stability at the lowlands is confined by uncertainty in thermodynamic data on minerals and atmospheric composition (Fig. 16.22), sulfatization of anorthite was confirmed to be more thermodynamically favorable at elevations where only Na-rich plagioclase could be stable. Here, we considered the stability of chlorides concerning atmospheric SO2 and showed that chlorides Ca, Mg, Na, and K are unstable concerning sulfatization (Fig. 16.24), exemplified by the net reaction  \n2NaCl + H2O(g) + 1.5SO2 (g) → Na 2SO4 + 2HCl(g) + 0.25S2(g). (16.73)  \nIn summary, multiple calculations of chemical equilibria at Venus' surface conditions suggest stability of alkali feldspars, Mg- and Mn-rich silicates, and sulfatization of Ca-rich pyroxenes and carbonates at ~100 -200 ppmv SO2 measured in the lower atmosphere. The stability of rock-forming minerals with respect to sulfatization on Venus decreases in the sequence, microcline, albite, enstatite, forsterite, rhodonite, tephroite, magnesite, rhodochrosite, anorthite, diopside, dolomite, wollastonite, calcite, natrite, and potassium carbonate. If near-surface f O2, f SO2, and f S2 are at conditions determined by the Mt-Py equilibrium (16.67) (Table 16.11), minerals less stable than anorthite should be subjected to sulfatization (Fig. 16.22). Similar conclusions can be made if near-surface f SO2 corresponds to 100 -200 ppmv SO2 and f O2 is in the range of 10 -20 -10 -21.7, estimated by Fegley et at. (1997b) for the conditions of modal radius. Although diopside is unstable, conditions of the diopside-anhydrite-enstatite-quartz equilibrium  \n<!-- image -->  \nCaMgSi2O6 + 1.5SO2(g) = CaSO4 + MgSiO3 + SiO2 + 0.25S2(g) (16.74)  \nare close to temperaturef (gas) terms suggested for the near-surface atmosphere (Tables 16.1, 16.2). Therefore, the sulfatization of diopside could be hindered, which is consistent with some experiments listed below.", 'Inferences from Laboratory Experiments': "In addition to many fundamental and technology-oriented studies of SO2-solid reactions, the alteration of carbonates, silicates, and silicate glasses has been investigated in the context of volcanic and postmagmatic processes (Burnett et al. 1997; Li et al. 2010; Auris et al. 2013; Henley et al. 2015; King et al. 2018; Palm et al. 2018; Delmelle et al. 2018; Renggli and King 2018; Renggli et al. 2018, 2019), and explicitly to constrain chemical weathering on Venus (Tables 16.9, 16.10). The venusian environments are challenging to replicate and monitor fully in the laboratory. Most Venus-focused experiments report insufficiently known partial pressures of p SO2 and other gases (S2, CO, OCS) that would affect the alteration. Experiments conducted in pressure vessels often cannot be held at pressure for long durations, thus leaving them to provide information only on the initial stages of alteration of geological materials. This contrasts with fundamental studies common in physical-chemical and chemical engineering literature, as Venus-focused experiments typically need to provide more information to assess rates and mechanisms of advanced sulfatization.  \nKinetic studies of SO2-calcite interactions in O2-bearing environments (sulfation) with applications to sequestration of SO2-rich industrial gases at ~700 -1200 K (e.g., Tullin and Ljungstróm 1989; Iisa et al. 1992; Tullin et al. 1993; Qiu and Lindqvist 2000; Hu et al. 2007; Jeong et al. 2015) provide information on the reaction rate and mechanism. The time-limiting step in the reaction mechanism involves the formation of a CaSO3 intermediate that is oxidized to CaSO4 by O2. The rate of SO2-CaCO3 reaction decreases with increasing f CO2 . Low f CO2 and high temperature favor the calcination of CaCO3 to CaO, which then reacts with SO2 gas. Meanwhile, elevated f CO2 at lower temperatures allows slower but fruitful sulfatization of calcite without calcination (e.g., Tulin et al. 1993). Solid-state diffusion of Ca 2+ toward the calcite surface and inward diffusion of SO2 through a porous CaSO4 have been considered to model reaction kinetics. Although sulfation experiments and kinetic models may not be directly applied to anoxic reactions (16.69) and (16.70), a rate-limiting step (16.69) of CaSO3 formation in both O2-rich  \n87  \n<!-- image -->  \nand anoxic conditions suggests the usefulness of rates and activation energies of sulfation reactions to assess sulfatization on Venus.  \nAlthough kinetic data on sulfation do not imply any long-term existence of calcite on the surface of Venus, a rapid anoxic sulfatization of calcite at the simulated surface environments has been confirmed. Fegley and Prinn (1989) investigated calcite alteration with a 1 bar CO2-SO2 mixture characterized by Venus-like p SO2 at 873 -1123 K. A calcite crystal was covered by ~10 µm thick anhydrite layer at 1123 K within eight days.  Although sulfatization of calcite at 90 bars CO2 could be slower than at 1 bar, the reaction rate (1 µm/yr CaSO4) assessed via extrapolation of results to ~740 K suggested geologically rapid alteration of calcite if it appears on the surface. Based on these values, Fegley and Treiman (1992a) noted that depletion of current atmospheric SO2 to f SO2 at the CaSO4-CaCO3 equilibrium (16.72) could occur in ~1.9 Ma. Aveline et al. (2011) observed sulfatization of calcite exposed to 4.8 bars of pure SO2(g) at 733 K. Sulfatization of calcite had been observed in several experiments at the NASA Glenn Extreme Environments Rig (GEER) facility at temperature, pressure, and simulated gas compositions of Venus' surface. Radoman -Shaw (2019) and Santos et al. (2023) reported the formation of 0.25 µm and 3 -5 µm thick coating of calcium sulfate in the GEER during 42 and 30 days, respectively. The discrepancy in results could reflect different degrees of SO2 consumption in reactions with stainless steel chamber walls, which has been proposed to occur in GEER experiments (e.g., Lucko et al. 2018). Slower sulfatization suggests more SO2 consumption and/or a more reducing atmosphere in which reactions (16.70) and (16.71) are suppressed at an elevated S2/SO2 ratio. Compositional mapping across a calcite alteration rind obtained by Santos et al. (2023) does not indicate sulfur diffusion into the calcite sample but suggests Ca 2+ diffusion toward the surface. Neither reaction rates nor mechanism of calcite alteration have been inferred from experiments that closely simulate Venus' surface conditions.  \nVenus-relevant experiments (Tables 16.9, 16.10) reveal slower reactions of SO2 with Ca-rich silicates and silicate glasses compared to calcite, and relative rates of sulfatization of silicates appeared roughly consistent with evaluations of mineral stability. Treiman and Fegley (1992a) confirmed a slower sulfatization of diopside than calcite based on experiments by Fegley and Prinn (1989). Slower sulfatization of Ca-silicates (wollastonite, diopside, augite, tremolite, etc.) was reported by Aveline et al. (2011), Radoman-Shaw (2019), Santos et al. (2023), and Port et al. (2023). In a 42-day GEER experiment, Radoman-Shaw (2019) observed widespread sulfatization of Ca-bearing pyroxenes, resulting in anhydrite formation and the appearance of sodium sulfates on aegirine and jadeite grains. Berger et al.  \n<!-- image -->  \n(2019) did not report an alteration of pyroxenes and plagioclase in basalt but observed some oxidation of olivine to iron oxide(s). Reid (2021), Reid et al. (2024), and Santos et al. (2023) did not indicate the sulfatization of plagioclases and pyroxenes at supposedly Venus-like p SO2. However, mineral phases were not analyzed explicitly in either set of experiments, and phase changes cannot be dismissed entirely. In experiments performed by Aveline et al. (2011), Radoman-Shaw (2019), and Santos et al. (2023), plagioclase feldspar appeared to be the most SO2-resistant Ca-bearing silicate, and no sulfatization was observed for thermodynamically stable phases such as Mg-rich silicates and alkali feldspars.  \nSulfatization of silicate glasses was observed in Venus-focused experiments at different temperatures, pressures, gas compositions, and durations (Berger et al. 2019; Reid 2021; Reid et al. 2024; RadomanShaw 2019; Radoman-Shaw et al. 2022; Santos et al. 2023). Sulfatization of Ca and Na in glasses occurred significantly faster than oxidation and/or sulfidation of ferrous iron, which has not always been observed. A more prominent sulfatization than oxidation was also observed in high-temperature experiments at elevated p SO2 irrelevant to Venus' environments (Cooper et al. 1996; Renggli and King 2018). Different degrees of alteration were observed in studies with compositionally similar MORB (Santos et al. 2003) and tholeiitic glasses (Reid 2021; Reid et al. 2024; Radowan-Shaw et al. 2022). Reid (2021) and Reid et al. (2024) reported a greater degree of alteration in alkaline basaltic glass than tholeiitic glass after a set of similar experiments was run using each material. Alteration of glasses by CO2-SO2 mixtures to sulfates and iron oxides occurred significantly faster than by CO2 (Teffeteller et al. 2022), which caused the oxidation of ferrous iron and carbonation of glass surfaces.  \nThe experiments of glass interaction with SO2-bearing gases led to the formation of sulfates of Ca and Na (anhydrite, thernadite, glauberite), together with less prominent iron sulfides and/or oxides. Different proportions of anhydrite, Na- and Na-Ca sulfates were observed in alteration products of compositionally variable basaltic glasses. Calcium sulfates were detected on the surface of basaltic glasses during a 42-day GEER run (Radoman-Shaw et al. 2022) and on a polished surface of MORB glass after a 30-day experiment in GEER (Santos et al. 2023).  Sodium sulfate dominated the alteration of alkaline glass products (Berger et al. 2019; Reid 2021; Reid et al. 2024; Radoman-Shaw et al. 2022). Reid (2021) and Reid et al. (2024) observed the formation of abundant sodium sulfate via the alteration of alkaline basalt glass and the appearance of anhydrite on tholeiitic glass surfaces during 14- and 30-day runs. Radoman-Shaw et al. (2022) reported the formation of abundant thernadite on a synthetic Venera  \n<!-- image -->  \n13-like K-rich glass with added sodium oxide. Contrary to high-temperature experiments conducted at high f SO2 (Renggli and King 2018), the formation of MgSO4 has not been observed in experiments relevant to Venus. This observation aligns with equilibrium thermodynamic assessments (Fig. 16.22).  \nThe observed formation of sulfate crystals on glass surfaces affects its texture. It alters the composition of subsurface layers that are affected by the migration of cations (e.g., Dyar et al. 2021). Compositional profiles through alteration rinds of alkaline glasses (Berger et al. 2019; McCanta et al. 2023; Reid et al. 2024) suggested migration of Ca and Na toward the surface to newly formed sulfates and relative enrichments in Si and Al in a subsurface layer. Berger et al. (2019) and Reid (2021) noted that the more prominent alteration of alkaline glasses likely reflects faster diffusion of Na than Ca, as known for silicate glasses (Behrens 1992). These inferences are mainly consistent with experiments on interactions of silicate glasses and minerals with SO2-rich high-temperature (> 1000 K) gases that suggest sulfatization through solid-state diffusion of cations (Ca, Na, and Mg), inward migration of electron holes and/or oxygen atoms, and modification of the Si-Al-O framework below the surface without formation of secondary Al-Si rich phases at initial stages of sulfatization (Auris et al. 2013; Henley et al. 2015; Matters et al. 2017; Delmelle et al. 2018; Palm et al. 2018; Renggli and King 2018; Renggli et al. 2019a, 2019b). Even in SO2-rich gas experiments, the diffusion of cations through sulfate coatings surpasses that of the inward diffusion of SO2, resulting in the formation of new sulfate on the surface of the coatings (Renggli et al. 2019a). Therefore, diffusion of cations from the substrate to the surface is likely the rate-limiting factor of sulfatization, at least in glasses (e.g., Auris et al. 2013; Renggli et al. 2019a).  \nIn experiments simulating venusian conditions, silicates and glasses typically remain partially uncoated by secondary phases due to the brief duration (< 80 days), except for alkaline samples. Relative sulfatization rates are inferred primarily from the early alteration stages. In addition to longer runs, further evaluation of rates and mechanisms of sulfatization at Venus' environments may require gas -solid experiments at higher temperatures and/or higher p SO2 and a thoughtful extrapolation of results to V enus' values. One possible way is to perform 1 bar runs at SO 2 molecular number density corresponding to Venus, as Fegley and Prinn did (1989). Although sulfatization rates at advanced stages of alteration are unknown from available experiments (Table 16.9), the age of basaltic materials on venusian plains (~0.3 -0.9 Ga, Herrick et al. 2023, for review) suggests a prominent sulfatization of exposed glasses and Ca-rich pyroxenes, consistent with the composition of surface materials (Table 16.3).  \n<!-- image -->", 'Discussion and Summary on Sulfatization': "The analyses of mineral stability and experimental results indicate the likelihood of sulfatization of Caand N-bearing carbonates, silicates, and basaltic glasses at the present surface. These materials exhibit different thermodynamic affinities for sulfatization and alter at varying rates. Several Ca-bearing minerals (e.g., plagioclase, F- and Cl-apatite, titanite) are thermodynamically stable, and/or their alteration is kinetically impeded. The chemical compositions of solids at the landing sites (Table 16.3) compared to calculations of corresponding equilibrium mineral assemblages (Table 16.8) suggest incomplete and uneven sulfatization of the sampled materials. Considering the exposure of sampled materials over several hundred Ma, the incomplete sulfatization is inconsistent with the experimental alteration of Ca-bearing phases and basaltic glasses (Table 16.9). The inconsistency could be attributed to slower alteration at advanced stages, to the resistance of certain phases to sulfatization, and/or to chemical equilibria between some minerals (e.g., labradorite, Barsukov et al. 1986c; Klose et al. 1992; Zolotov 2018), their sulfatization products, and gases. As an example, the following chemical equilibrium could impede the alteration of plagioclase in basalts,  \nNa0.4Ca0.6Al1.6Si2.4O8 (labradorite) + 0.9SO2(g) ↔ 0.6CaSO4 + 0.4NaAlSi3O8 + 0.6SiO2 + 0.6Al2SiO5 + 0.15S2(g). (16.75)  \nA complete gas-solid equilibration of exposed materials, as modeled by Barsukov et al. (1982c, 1986c), Klose et al. (1992), and Semprich et al. (2020), is not supported by the variable sulfur content and the S/Ca ratio in the samples (Tables 16.3, 16.8). The different degrees of sulfur intake may reflect differences in exposure time, fugacities of reacting gases, surface area, grain size, surface/volume ratio in exposed materials, the abundance of glass, the plagioclase/pyroxene ratio, the amount of Na-bearing phases subjected to sulfatization, and the abundance of secondary sulfides. Interestingly, the abundance of sulfur (Table 16.3) does not correlate with potassium content or the Al/Ca ratio , which could reflect the plagioclase/pyroxene ratio and the iron content that could affect the degree of sulfidation. The advanced physical weathering (formation of smaller-size particles) of Venera 13 rocks (Fig. 16.1; Basilevsky et al. 1985; Garvin et al. 1984) compared to Venera 14 coincides with higher sulfur content in the former sample (Zolotov and Volkov 1992). However, faster alteration of alkaline glasses could have contributed to more sulfur intake by K-rich Venera 13 rock than by Venera 14 counterpart. The increasing amount of secondary sulfates in exposed basaltic materials with exposure age would affect  \n<!-- image -->  \nphysical properties (e.g., near-IR emissivity, Dyar et al. 2020, 2021; McCanta et al. 2024) that could be used to diminish freshly exposed materials such as lavas.  \nThe thermodynamic instability of Ca-rich plagioclase and the lesser stability of diopside, along with other Cabearing silicates and mafic glasses, at the conditions of the Venus' highlands (Zolotov 2018), could lead to more prominent sulfatization at elevations. Admixtures of anhydrite in crashed basalt samples (Dyar et al. 2021) and sulfate coatings of 0.1 mm have been found to substantially enhance the near-IR reflectance of basalt (McCanta et al. 2024). Such admixtures and coatings have the potential to create 'white' mountains. Putative sulfate coating could have formed gradually since the inception of the highland rocks. Alternatively, or additionally, they could be a consequence of intense sulfatization under conditions of elevated volcanic f SO2 and temperature during the formation of the volcanic plains ( Sect. 16.4.4.1 ). Regardless of the cause of possible high-altitude sulfatization, the low near-IR emissivity at highlands ( Sect. 16.2.2.4 ) could be caused by low-albedo sulfate-rich alteration materials of basalts and basaltic glasses.", '16.3.2.3  Stability and Reactivity of Sulfides': 'Despite analytical, numerical, and experimental studies, the fate of metal sulfides at the surface is less clear than that of sulfates. The ambiguity reflects the uncertainty of thermodynamic properties of minerals and their solid solutions, unknown abunda nces of SO₂, OCS, S₂, and CO below 12 km (Table 16.1), model-dependent constraints on the composition of the near-surface atmosphere, and the absence of proof of gas-phase equilibration at the lowlands. The lack of such equilibration above a thin atmospheric layer (Krasnopolsky and Pollack 1994; Zolotov 1996) restricts the evaluation of mineral stability at highlands (Zolotov 2018). It follows that f S2 and f O2 values, commonly used to assess the stability of coexisting sulfides and oxides, are vague at the conditions of modal planetary radius and may not be defined at altitudes in the framework of equilibrium chemical thermodynamics. For the Fe-S-O system (Figs. 16.22, 16.24, 16.25), even small changes in the atmospheric chemistry could affect the composition of stable phases. Since the 1970s, researchers have employed various atmospheric compositions and assumptions regarding gas-phase and gas-solid equilibria, often yielding inconsistent conclusions regarding sulfide stability.  \n<!-- image -->', 'Iron Sulfides and Oxides in Surface Materials and Buffering of Atmospheric f O2 and f S2': "As mentioned in Sect. 16.3.2.2 , Mueller (1965) first considered the stability of iron sulfides and oxides at presumed conditions of the venusian surface and discussed their influence on the composition of trace S- and C-bearing atmospheric gases. He has demonstrated that S-bearing gases should not be major atmospheric species in contact with iron sulfides at ~700 K. In his opinion, the measured abundances of such gases and the CO2/CO ratio will inform the mineralogy of sulfides and oxides in upper crustal materials. Without developing mineral stability diagrams for the Fe-O-S-H and Ca-Si-O-S-C systems, he proposed that the highest possible atmospheric CO2/CO ratio could be controlled by coexisted hematite and magnetite (Mueller 1964) via (16.65) that sets O2 fugacity of the near-surface atmosphere by (16.66). Mueller (1965) noted that such oxidizing conditions do not exclude a coexistence of pyrite with iron oxides but do not allow the stability of troilite and pyrrhotite. According to his viewpoint, the magnetite-pyrite-hematite (Mt-Py-Hem) equilibrium (a sum of (16.66) and (16.67)), could affect partial pressures of trace S-bearing atmospheric gases (SO2, OCS, SO3, Sn, etc.) and the oxidation state ( f O2, CO/CO2) of the gas. He also considered the magnetite-pyrite (Mt-Py) equilibrium (16.67), which could control S-bearing gases at lower f O2 than the magnetite-hematite (Mt-Hem) equilibrium (16.66). The concept of Mueller (1965) remains valid in 2025. The atmospheric composition ( Sect. 16.2.1 , Table 16.1) is consistent with Mueller's inferences about the instability of reduced ferrous sulfides and with coexisting iron oxides with pyrite that could constrain gas chemistry in pore spaces of permeable crustal materials and near-surface atmosphere. Because a near-surface gas composition sets boundary conditions for atmospheric models (e.g., Krasnopolsky and Parshev 1981; Krasnopolsky 2007, 2013; Bierson and Zhang 2020), it likely affects the whole lower atmosphere. Table 16.11 (col. 3) depicts the assessed atmospheric composition that corresponds to measured abundances of CO2, SO2, and H2O and trace gases (CO, OCS, Sn, H2S, etc.) at f O2 and f S2 controlled by the magnetite-pyrite equilibrium.  \nLewis (1968) considered gas-solid equilibria involving troilite, pyrite, magnetite, ferrous sulfate, olivine, and enstatite to constrain unknown atmospheric abundances of H2S, OCS, and SO2. The estimated mixing ratios of S-bearing gases (< 10⁻⁶) were below their observational upper limits, and reduced gases controlled by troilite-bearing mineral assemblages dominated over SO₂. As noted in Sect. 16.3.2.2 , Lewis (1970) estimated possible mineral-buffered atmospheric compositions at the f O2 set by either graphiteCO2-CO or magnetite-silicate equilibria. Assessed mixing ratios of OCS and H2S (50 ppmv and 5 ppmv at 748 K and 120 bars) in equilibria with troilite appeared more abundant than SO2 (0.3 ppmv) and SO3  \n<!-- image -->  \n(~10 -16 ) supposedly controlled by calcite-anhydrite equilibria in surface/near-surface materials ((16.64), (16.72)). Following the work of Mueller (1965), these estimations demonstrated that troilite in surface materials and low f O2 set low atmospheric f SO2 and f SO3, allowing the stability of Ca-rich carbonates and silicates. While the mineral stability diagrams created by Lewis and Kreimendahl (1980) reinforced this conclusion, the in situ data on the high SO2/(OCS + H2S) ratio in the sub-cloud atmosphere published later in 1980 (Table 16.1) suggested a more oxidizing atmosphere. This atmosphere permits the stability of pyrite ( Sect. 16.3.2.2 ) over troilite or pyrrhotite and favors iron oxides over ferrous silicates.  \nThe consideration of the PVLP and Venera 12 data on SO2, CO, and CO2 in calculations of gas-solid multisystems by Barsukov et al. (1980a, b, 1982c) and Khodakovsky (1982) inferred instability of troilite and pyrrhotite at the surface. Pyrite and/or magnetite appeared in their equilibrium mineral assemblages, depending on the gas composition, temperature, and pressure used in the models. Barsukov et al. (1980b, 1982c) inferred from these fluctuations that a pyrite-magnetite-anhydrite assemblage might be present in surface materials, potentially buffering the composition of the nearsurface atmosphere.  \nVon Zahn et al. (1983) and Prinn (1985) proposed that pyrite connects the atmospheric and geological sulfur cycles. They anticipated the consumption of crustal pyrite via thermal decomposition and interactions with atmospheric CO2 and H2O, resulting in the formation of S2, OCS, and H2S, respectively. In their schemes, net photochemistry-driven oxidation of reduced gases to SO2 was followed by sulfatization of exposed calcite to anhydrite. A conversion of anhydrite to pyrite and calcite through unspecified crustal processes was suggested to complete the geological sulfur cycle. However, the involvement of calcite disagreed with its compelling instability in contact with atmospheric SO2, CO, and S2 (Barsukov et al. 1980b, 1982c; Fegley and Prinn 1989; Zolotov 1985), and the reduction of anhydrite to pyrite had not been supported either numerically or experimentally. Fegley et al. (1995, 1997a) slightly modified the geologic sulfur cycle of von Zahn et al. (1983) based on their experiments on pyrite → pyrrhotite → magnetite conversion at simulated venusian conditions. In their pathways, pyrite from igneous rocks thermally converts to metastable pyrrhotite by reaction  \n(1 -x )FeS2 → Fe 1x S + (0.5 -x )S2(g). (16.76)  \n<!-- image -->  \nInteractions of pyrrhotite with CO2 and H2O produce magnetite and supply OCS and H2S to the atmosphere, and anhydrite is not reduced to complete the cycle. In other words, Fegley et al. (1995) proposed net oxidation of sulfides to sulfates in upper crustal materials that occurs via photochemistrydriven oxidation of Sn, OCS, and H2S to SO2. Although the stability and reactivity of pyrite remain questionable, this vision aligns with the current understanding of sulfur fate in the atmosphere-surface system (Fig. 16.15).  \nAlthough the instability of pyrrhotite and troilite at the surface has been evident since 1980, calculations of chemical equilibria have resulted in varied conclusions about the stability of pyrite at the surface. The stability of pyrite in the lowlands has been inferred from calculations of chemical equilibria in gas-basalt multi-systems (Barsukov et al. 1980a, b, 1982c; Klose et al. 1992) as well as from calculations of selected gas-mineral equilibria (Zolotov 1991a, b, c, 1992, 2018, 2019; Zolotov and Volkov 1992; Hashimoto and Abe 2005) such as (16.67) and  \nFe3O4 + 6SO2(g) + 16CO(g) = 3FeS2 + 16CO2(g). (16.77)  \nThese evaluations suggest the formation of secondary pyrite through the interaction of reduced Sbearing gases with ferrous glasses and minerals and via oxidation of pyrrhotite by CO2, H2O, SO2 (Zolotov 2018), and S2 (back reaction (16.76)),  \nFeO (in minerals and glasses) + 2COS(g) → FeS 2 + CO2(g) + CO(g) (16.78)  \nFeO (in minerals and glasses) + 2H2 S(g) → FeS 2 + H2O(g) + H2(g) (16.79)  \n2Fe0.875S + CO2 (g) → 0.25Fe 3O4 + FeS2 + CO(g) (16.80)  \n3Fe0.875S + 1.125S2 (g) → 2.625FeS 2 (16.81)  \nthough concrete pathways remain to be inferred experimentally.  \nExperimental data (e.g., Lv et al., 2015; Aracena and Jerez 2021; Liu et al. 2023) demonstrate that both thermal decomposition and oxidation of pyrite occur via the formation of pyrrhotite (Eq. (16.81)). With application to Venus, experiments by Treiman and Fegley (1991) and Fegley et al. (1995) revealed the alteration of pyrite in the presence of CO2-rich gas mixtures at 1 bar (Table 16.9). Fegley et al. (1995) reported a more rapid pyrite-to-pyrrhotite conversion (16.76) compared to the  \nsubsequent oxidation of pyrrhotite to stable iron oxides.  Wood and Brett (1997) expressed concerns about the relevance of the p S2, presumed but not measured in the experiments of Fegley et al. (1995), to the conditions on Venus. In his reply, Fegley (1997) argued for pyrite instability at supposedly low Venus' f S2 and observation-based f S3 (Krasnopolsky 1987) grounded on considerations within the Fe-S-O phase system, using coordinates of f S3, f S2, and f O2. Hong of Fegley (1997b) studied the kinetics of the pyrite-to-pyrrhotite conversion in various gas mixtures at 1 bar. They demonstrated the feasibility of the reaction at Venus' temperatures. Experimental data on f S2 over pyrite, as measured and reviewed by Hong and Fegley (1998), further supported the feasibility of (16.76) at 740 K and f S2 < 10 -5.0 bars. It follows that Venus' pyrite is unstable concerning thermal decomposition to pyrrhotite at f S2 of 10 -5.7 bars which corresponds to ~20 ppbv S2 at chemical equilibrium with 9-13 pptv S3 assessed by Krasnopolsky (2013) for 3-10 km altitude (Table 16.1; Model 10 of Table 16.2) and to the chemical equilibrium with 8 ppmv CO (Model 9 of Table 16.2). However, the instability of pyrite at that low f S2 does not necessarily indicate its weathering on Venus. This is because the mineral is not typically found in fresh basaltic materials that are more reduced than the QFM f O2 buffer, as inferred for the three Venus' surface probes ( Sect. 16.2.2.1 ).  \nReid (2019) and Reid et al. (2024) used pyrite to introduce S-bearing gases in their basalt-weathering experiments in which the magnetite-hematite buffer was used to set the oxidation state ( f O2, CO2/CO, SO2/S2, etc.) of the reaction vessel gas. Based on the experimental work of Lv et al. (2015), they expected SO2 to dominate among S-bearing gases produced via pyrite-CO2 interaction. However, pyrrhotite, hematite, and magnetite in post-run assemblages indicated incomplete pyrite oxidation. These four minerals cannot stably coexist (Figs. 16.22, 16.24, 16.25), and the occurrence of these phases limits a proper assessment of the gas atmosphere during the runs. One possible explanation of pyrrhotite occurrence is a partial reduction of gases (CO2 → CO, OCS; SO2 → OCS, S2) on the steel walls of the reaction vessel. Such a reduction may also account for the limited oxidation of basalts in these and other experiments (Berger et al. 2019; Radoman-Shaw et al. 2022; Santos et al. 2003) in steel vessels.  Although the observed pyrite-to-pyrrhotite conversion is consistent with preceding works (Treiman and Fegley 1991; Fegley et al. 1995; Hong of Fegley 1997b; Lv et al. 2015), the results of Reid with colleagues may not constrain the fate of pyrite at Venus' lowlands.  \nIf Venus' surface f S2 is 10 -4.4 -10 -4.7 bars (~0.2-0.4 ppmv S2 at 95.6 bars total pressure) inferred from the supposed equilibration of major atmospheric gases (Models 5, 7, and 8 of Table 16.2), stable pyrite  \n<!-- image -->  \ncould coexist with magnetite (Figs. 16.22, 16.25; Table 16.11). Berger et al. (2019) reported the formation of pyrite in experiments on basalt alteration in a low-temperature run at Venus' gas pressures. Santos et al. (2024) observed unaltered pyrite and Fe-sulfide-to-pyrite conversion in a 60-day GEER experiment at 93 bars and 733 K. Radoman-Shaw (2019) noted oxidation of pyrrhotite and pyrite to magnetite in GEER runs. Again, we note that the use of stainless-steel chambers and vessels in most Venus-relevant experiments to address the stability of sulfides (e.g., Berger et al. 2019; Kohler 2016; Port et al. 2016, 2018; Port and Chevrier 2017a, b, 2020; Radoman-Shaw 2019) could have altered the composition and redox state of reactive gases. The restricted capability to monitor or regulate the gas composition in these experiments complicates the application of these findings to Venus. Regardless of the initial presence of pyrite and pyrrhotite in exposed surface materials, their chemical pathways need accurate and meticulous experimental investigations under controlled gas composition.  \nThe stability fields of solids in the Fe-O-S system indicate that environments at Venus' lowlands closely resemble the conditions under which pyrite, magnetite, and hematite coexist (Klose et al. 1992; Semprich et al. 2020; Zolotov 1991a, b, c, 1992b, 2018, 2019). Those phase diagrams demonstrated that a definitive conclusion regarding mineral stability cannot be drawn without precise information on the atmospheric composition at corresponding temperature, pressure, and altitude. Stability diagrams created by Fegley (1997) and Zolotov (1991b, c, 1992, 2018, 2019) demonstrated that varying conclusions on sulfide stability were drawn based on the use of different instrumental data for atmospheric composition (S3, SO2, OCS, H2S, and CO) and a range of assumptions on gas-phase equilibration. This is further illustrated in Fig. 16.25. Ambiguity also mirrors the uncertainty of the thermodynamic data used on solid phases, leading to imprecise conditions of phase equilibria. This is evident in the Mt-Hem equilibrium (refer to (16.65) and (16.66), Figs. 16.22 and 16.25, Fegley et al. 1997b) and the Py-Pyrr boundary, which varies in the works of Hong and Fegley (1998), Semprich et al. (2020), and Zolotov (1992b, 2018, 2019).  \nTemperature, pressure, and suggested gas composition at the modal radius are close to the conditions determined by the magnetite-pyrite assemblage suggested elsewhere (Mueller 1965; Barsukov et al. 1980b, 1982c; Klose et al. 1992; Hashimoto and Abe 2005; Wood 1997; Zolotov 1992b; Zolotov and Khodakovsky 1989; Zolotov and Volkov 1992). An assemblage like this in permeable surface materials could control ratios between oxidized and reduced gases (e.g., SO2/OCS/S2) in the near-surface  \n<!-- image -->  \natmosphere and shallow subsurface, provided those chemical equilibria (16.67) and/or (16.77) are achieved. The Mt-Py-Pyrr and Mt-Py-Hem invariant equilibria bracket these ratios  \n0.375Fe3O4 + 1.5FeS2 = 3Fe0.875S + 0.75O2 (16.82)  \nFe2O3 + 0.25FeS2 = 0.75Fe3O4 + 0.25S2 (16.83)  \nthat set the H-C-O-S system's fugacity of S2, O2, and other gases (Figs. 16.22, 16.24, 16.25, Table 16.11). For example, the Mt-Py-Hem equilibrium could control f SO2,  \n16Fe2O3 + FeS2 = 11Fe3O4 + 2SO2 (16.84)  \nThe considerable uncertainty of f O2 set by the Mt-Hem equilibrium (16.66) (Fegley et al. 1997b) does not exclude the existence of stable hematite together with magnetite and pyrite at the surface of Venus. If this is the case, the Mt-Py-Hem assemblage in surface materials could control f S2 (< 10 -4.17 bars, (16.83)), f O2 (< 10 -20.8 bars, (16.66)), and the composition of other gases. The composition of such gases (e.g., x SO2 < 713 ppmv, x CO > 9 ppmv) is more consistent with measurements (Table 16.1) than that calculated in Table 16.11 for the nominal Mt-Py-Hem buffer.  Thermodynamic calculations of Zolotov (1994, 1995b) and Semprich et al. (2020) demonstrated that substituting TiO2 in hematite, forming a hematite-ilmenite solid solution, results in lower f O2 at the Mt-Hem equilibrium. This outcome aligns more closely with near-surface atmospheric models (Table 16.2), constrained by measurements. The presence of hematite in widespread coatings and fine-grained materials is consistent with near-infrared reflectance at Venera landing sites (Pieters et al. 1986; Shkuratov et al. 1987, Sect. 16.2.2.1 ).  \nIn addition to chemical weathering, non-silicate iron phases could be delivered from clouds. Zolotov (2021) suggested the accumulation of hematite in surface materials through chemical and/or thermal decomposition of unstable ferric sulfate (Fig. 16.25) that forms from cosmic iron in the clouds and sinks to the surface in particles. In addition to in situ data on iron in clouds, ferric sulfates could be UV-blue cloud absorbers ( Sect. 16.3.1.8 ). Ferric sulfate could have been captured at the PVLP LNMS inlet in lower clouds and thermally decomposed to hematite and gaseous sulfur oxides, as inferred in the instrument's mass spectra (Mogul et al. 2024; Zolotov et al. 2023).  \nAlthough pyrite and iron oxides are expected secondary minerals on lowlands, the low dielectric constant of lowland materials estimated from the Magellan radar data and ground-based radiometric  \n<!-- image -->  \ndata ( Sect. 16.2.2.1 ) does not indicate abundant iron sulfides and magnetite formed through gas-solid weathering reactions in present epoch in upper (< 0.5 m) surface layer. The supposed presence of more abundant iron sulfides and/or oxides in deep (>1 km, Pettengill et al. 1992) and shallow (>0.7 m, Anotony et al. 2022) subsurface suggested from radar data could reflect the mineral formation formed at preceding climatic conditions followed by burying by lava and fine-grained materials ( Sect. 16.4.4 ).  \nIf the near-surface abundance of S-O-C-H gases is governed by the Mt-Py and/or Mt-Py-Hem assemblages (Figs. 16.22, 16.25, Table 16.11), the role of anhydrite in buffering remains ambiguous. Despite the limited sulfatization at the landing sites (Table 16.8), the experimentally noted reactivity of basaltic glasses with SO2 (Tables 16.9 and 16.10) and the thermodynamic favorability of sulfatization for Ca-bearing pyroxenes, there is no clear indication of equilibration for essential sulfatization reactions ( Sect. 16.3.2.2 ). This does not necessarily exclude the participation of anhydrite in the buffering process. Indeed, kinetic inhibition of advanced sulfatization and/or the equilibration of plagioclase with its weathering products (Eqs. 16.71 and 16.75), as suggested by equilibrium models (Barsukov et al. 1986c; Klose et al. 1992; Zolotov 2018), could decelerate the depletion of f SO2 towards values at which CaSO4bearing mineral assemblages are in equilibrium (e.g., (16.75)). The participation of anhydrite in buffering by the Mt-Py assemblage suggested by Barsukov et al. (1982c) could be illustrated by equilibrium from Fegley and Treiman (1992a)  \n3FeS2 + 6CaAl2Si2O8 + 11O2 = 6CaSO4 + Fe3O4 + 6Al2SiO5 + 6SiO2 (16.85)  \nthat sets f O2 of 10 -21.5 ± 0.4 bars at 740 K. Current atmospheric x SO2 of ~150 ppmv could reflect a slow ongoing sulfatization of Ca-bearing pyroxenes, while corresponding f SO2 of 10 -1.84 bars could control concentrations and fugacities of trace S-bearing gases at the Mt-Py equilibrium at the surfaceatmosphere interface (Table 16.11; Fig. 16.22). The buffering of atmospheric gases by secondary minerals on Venus' lowlands may have resulted from a coupled chemical evolution of the atmosphere -surface system since the last global volcanic resurfacing event ( Sect. 14.4.4 ).", 'Fe-O-S Minerals at Highlands': 'The lack of equilibration among gases above a thin near-surface layer restricts evaluations of phase equilibria in the highlands, and such assessments heavily rely on assumptions. If gas-phase equilibria are assumed at the highlands, decreasing f O₂ with altitude allows the stability of pyrite. Indeed, calculations  \n<!-- image -->  \nof complete equilibria in multicomponent gas-solid systems indicate the stability of pyrite under highland conditions (Khodakovsky 1982; Klose et al. 1992; Semprich et al. 2020). If f O₂ is not affected by gas equilibria and remains the same as at the lowlands, hematite is stable at the highlands (Fegley 1997; Fegley et al. 1997b; Semprich et al. 2020; Zolotov 1987, 1996; Zolotov and Khodakovsky 1989). However, when evaluating the stability of iron sulfides and oxides, the results drastically depend on the chosen gas-solid chemical equilibria and the measured abundances of gases. Phase diagrams, plotted in terms of f S2, f S3, f SO2 , f OCS , f CO, and f H2S in the works mentioned above, do not rule out the presence of hematite, magnetite, pyrite, or even pyrrhotite in the highlands. The confusing discrepancy arises from the deviation of extrapolated measured gas abundances (Table 16.1) from the conditions of gaseous equilibria with increasing altitude above a narrow near surface layer at the modal radius (Zolotov 1996). Consequently, chemical equilibrium approaches may not be directly applicable to assess the reactivity of solids exposed to the highland atmosphere. Regardless of further remote sensing or in situ data on surface materials, experimental studies on the kinetics of gas-mineral reactions with selected gases or disequilibrium gaseous mixtures are necessary to understand the secondary mineralogy of the highlands, some of which could have formed during the global volcanic resurfacing ( Sect. 16.4.4 ).', 'Chalcogenides of Trace Elements in Surface Materials': "The stability of trace metals sulfides in surface and upper crustal materials have been examined in the  \ncontext of cloud composition ( Sect. 16.2.1.4 ), low microwave emissivity of highland materials ( Sect. 16.2.2.5 ), and the formation of hazes in the lower atmosphere. Lewis (1968) was the first to consider the migration of metal-bearing compounds via sublimation from rocks and condensation on the presumed cold polar regions on Venus. Lewis (1968, 1969) investigated the possible condensation of sulfides and chalcogenides of mercury in the clouds. He concluded that the high surface temperature and a limited abundance of atmospheric mercury prevent the condensation of HgS (cinnabar) anywhere on the surface. Similarly, Brackett et al. (1995) considered the vapor pressures of metal chalcogenides and halogenides as a function of temperature to determine the likelihood of their sublimation from lowland materials and condensation on highlands and in the lower atmosphere below 15 km. For chalcogenides, they observed a decrease in vapor pressure at 740 K in the following sequence: As2S2, CuS, HgS, NiS2, Sb2S3, CoS2, SnS, PbS, Ag2S, Bi2S3, and ZnS. Brackett et al. (1995) demonstrated the potential for sublimation and condensation of the compounds considered above and assessed the rates of sublimation, eddy transport, and condensation of solids. They concluded that a 0.1 -1 cm thick layer  \n<!-- image -->  \nof metal chalcogenides and chlorides could accumulate on highlands throughout 1 -10 Ma. Brackett et al. (1995) noted that the high dielectric constants (180, 190) of potentially condensable Sb2S3 (stibnite) and PbS (galena) might contribute to the low microwave emissivity at highlands. Their work suggests that condensation could contribute to the formation of near-surface (< 6 km) atmospheric hazes tentatively suggested from in situ observations (Ragent and Blamont 1980; Gregier et al. 2004).  \nSchaefer and Fegley (2004) further investigated the potential condensation of volatile species as a function of altitude from the surface to ~51 km. They calculated solid-gas equilibria constrained by estimated partial pressures of volatile compounds of trace elements (Hg, Zn, Cd, Cu, Pb, Ag, Zn, Sb, Bi, As, etc.) that could be sublimated or volcanically outgassed from the crust. They used element/sulfur ratios in the terrestrial oceanic crust as a proxy for Venus' crust. They assumed that the atmospheric sulfur abundance (Table 16.12) represents the degassing degree of other elements. Schaefer and Fegley (2004) predicted condensation of As2S3, Sb2S3, and Bi2S3 (bismuthite) above 27, 17, and 1.6 km, respectively, and condensation PbS, Cu2S, ZnS, Ag2S, CdS at all hypsometric levels. PbS and Bi2S3 could condense above ~2.6 km if atmospheric abundances of Pb and Bi correspond to depleted crustal abundance and/or inefficient sublimation. In such a case, both sulfides could contribute to the low microwave emissivity of the highlands. Schaefer and Fegley (2004) also discussed possible condensation of lead-bismuth sulfosalts (e.g., PbBiS4, Pb3BiS6), known as fumarolic minerals. Port and Chevrier (2021) used the same approach to investigate the gas-solid equilibria for HgS and HgTe. They showed that condensation in the highlands is improbable due to insufficient Hg and Te in crustal rocks available for sublimation.  \nInferences about degassing and condensation of trace metal compounds by Lewis (1969), Schaefer and Fegley (2004), and Port and Chevrier (2021) are constrained by the unknown atmospheric compositions, with assessed gas abundances likely representing upper limits. Additionally, an overwhelming occurrence of crustal trace elements in dissimilated rather than mineral form (e.g., PbS, HgS) would have limited their degassing through sublimation. The apparent lack of hydrothermal processes, which can concentrate trace metals, in upper crustal materials also decreases the likelihood of an ample supply of trace element sulfide minerals that can sublimate. Lastly, the release of sulfur into the atmosphere from pyrrhotite oxidation (e.g., (16.80) and (16.81)) affects the ability to use atmospheric sulfur as an anchor in the estimations mentioned above.  \n<!-- image -->  \nSeveral qualitative exploratory experiments have explored the behavior of trace sulfur-bearing compounds in the atmosphere-surface system (Table 16.9). Kohler et al. (2013) and Kohler (2016) demonstrated the sublimation of HgS under simulated conditions resembling Venus' mountaintops (653 K, 55 bars CO2) within two days. Port and Chevrier (2017a) and Port et al. (2018) noted the complete loss of HgS in the 1-bar CO2 atmosphere at 653 -733 K in 1 day, and lesser losses were observed in CO2-SO2 and CO2-OCS gas mixtures. They observed oxidation of galena to PbSO4 (anglesite) in all runs, and Port et al. (2019) reported stable anglesite in these gas mixtures at Venus' temperatures and pressures. Port et al. (2020) investigated the behavior of Te-Bi-S mixtures and Bi2S3 at lowland and highland conditions in the presence of CO2, CO2-SO2, and CO2-OCS gas combinations. They reported a preferential formation of Bi2Te2S (tetradymite), and additional phases (e.g., Bi2S3, Bi2Te3) appeared depending on the experimental conditions.", '16.4   Sulfur in the History of Venus': "Sulfur on Venus could have played significant roles in differentiation and core formation, in the magma ocean, in magmatic degassing, in atmospheric processes and climate, in trapping of degassed volatiles to minerals, and in exchanges of crustal and mantle materials via igneous, tectonic, and putative aqueous processes. The following questions remain unanswered, and the ambiguity limits the understanding of Venus' sulfur. How much water did Venus accrete during its formation? Did Venus lose water from its atmosphere, and if so, how and when? Did hydrogen escape lead to oxidation of atmospheric, crustal, and mantle materials? Did water condense on the surface to form oceans? Did crustal materials subduct into the mantle? Did Venus have life? Understanding Ve nus' evolution is limited due to the lack of solid samples and because radar images provide insights into only about 1/10 to 1/5 of Venus' history. Although geological, climatic, and coupled atmospheric-interior evolution has been addressed in multiple works (see Avice et al. 2022; Herrick et al. 2023; Ghail et al. 2024; Gillmann et al. 2022, O'Rourke et al. 2023, Rolf et al. 2022, and Salvador et al. 2023, for reviews), the behavior of sulfur has not been detailed. The behavior of sulfur remains a mystery, and its detailed role in Venus' evolution requires investigation. This section briefly describes current views on Venus' evolution, provides insights into the fate of sulfur, and addresses publications in which sulfur is considered.  \n<!-- image -->", '16.4.1 Formation and Early History': "Accretion models suggest the formation of Earth and Venus from compositionally similar materials of the protoplanetary disk (e.g., Wetherill 1978; Hansen 2009; Morbidelli et al. 2012; Walsh and Levison 2016; Raymond and Izidoro 2017; Raymond et al. 2020). Although accretion from planetesimals and planetary embryos is a commonly accepted pathway (Raymond et al. 2020; Raymond and Morbidelli 2022), accretion from pebble-size materials is not excluded (Johansen et al. 2021). Although the stochastic accretion of planetary embryos and planetesimals (e.g., Raymond et al. 2009; Sossi et al. 2022) suggests compositional differences between the planets, drastic variances are not expected, and pebble accretion could provide more similar compositions. The compositional similarity of the planetbuilding materials of Earth and Venus is consistent with the comparable planetary sizes, densities, moments of inertia, masses of carbon and nitrogen in the combined atmosphere-crustal materials (Table 16.12) and basaltic materials with similar composition and oxidation state ( Sect. 16.2.2 ).  \nNotable compositional differences between the planets, such as atmospheric and surface water content, could reflect divergent evolution patterns rather than the initial ones. The higher abundances of 20 Ne and 36 Ar in Venus' atmosphere could reflect the preservation of the primordial atmosphere due to fewer early giant impact(s) that could have stripped an early terrestrial atmosphere (Genda and Abe 2005; Sakuraba et al. 2019). Venus' water deficiency is commonly attributed to the photodissociation of water vapor and the following hydrogen escape. The escape is consistent with the anomalously high atmospheric D/H ratio of (2.4 ± 0.5) × 10 -2 (Donahue et al. 1997), though the amount of accreted water may not be directly inferred from these data (e.g., Grinspoon 1993; Greenspoon and Lewis 1988; Zahnle and Kasting 2023).  \nPossible accretion of Earth and Venus from Mars-mass differentiated planetary embryos (Morbidelli et al. 2012; Raymond et al. 2020; Raymond and Morbidelli 2022; Izidoro et al. 2021) suggests the delivery of sulfur that was partitioned into cores of those bodies as sulfide. Although troilite is likely to be sulfide in cores, its abundance and occurrence could vary depending on bulk composition with the body and its differentiation pathways. Differentiated bodies with the composition of ordinary chondrites could have relatively troilite-poor metal cores, and differentiated carbonaceous chondritic bodies could have pure troilite cores (Bercovici et al. 2019) that remain to be proved in iron meteorite samples, however. Although iron meteorites (many of which are nuggets of planetesimal cores) commonly contain only  \n<!-- image -->  \n~0.5 -1 wt% S (Goldstein et al. 2009), concentrations of siderophile elements in iron meteorites suggest an initial ~0 -19 wt% S in the cores of asteroids and planetesimals (Chabot and Zhang 2022; Tornabene et al. 2023; Zhang et al. 2024). Some 'missing' sulfur could be lost via collision-induced volatilization of sulfides (Rubin et al. 2022), consistent with an apparent depletion in other moderately volatile elements (Cu, Ga, and Ge) in iron meteorite samples, or attributed to lower mechanical resistance of sulfides to attrition in space, during atmospheric passage, or weathering on Earth (Kracher and Wasson 1982). Alternatively, some sulfur (in sulfides) could be in the core's parts that are underrepresented in the iron meteorite collection (Ni et al. 2020). Sulfides could be in putative inner cores, troilite core-sourced lavas, or mantle intrusions formed during the crystallization of outer metal cores (Bercovici et al. 2019; ElkinsTanton et al. 2020). Preserved core's sulfides are a hypothesis, and the possible accretion of Venus and Earth from S-depleted differentiated non-carbonaceous planetesimals does not imply the formation of S-rich planetary cores. A low sulfur content in the solid inner core on Earth in the presence of Si in the outer core is supported by the experiments of Sakai et al. (2023) on the partitioning of Si and S between solid and liquid Fe under core pressures. Their results show that the Earth's core includes O and/or H, cannot be Fe-S-Si alloys, and contains 0.7 -1.7 wt% S. Low core sulfur contents have been estimated at 0.6 wt% and 1.9 wt% by Dreibus and Palme (1996) and McDonough (2014), respectively. Models of Earth's composition, based on chondritic components by Alexander (2022), suggest sulfur contents ranging from 0.70 to 0.82 wt% S.  \nThe Moon is formed by a giant collision with differentiated Earth ~4.47 Ga or ~100 Ma after forming the Ca-Al-rich inclusions (CAIs) (Canup 2012). The possible lack of the Moon-forming collision on Venus could have contributed to differences among planets regarding core formation from embryos' cores, and Venus could have experienced lesser largescale melting. However, Venus' slow retrograde rotation does not exclude a giant impact that altered the angular momentum (and spin direction) but did not lead to the formation of a moon (Canup et al. 2001; Davies 2008). Jacobson et al. (2017b) did not exclude an early (~11 Ma after CAI) giant impact on Venus followed by less energetic collisions that did not disrupt a stratified core, as discussed below. If a significant fraction of undifferentiated planetesimals and pebbles contributed to planet formation (Johansen et al. 2021; Brož et al. 2021), the delayed segregation of metal-sulfide cores involved the melting of silicates driven by the radioactive decay.  \n<!-- image -->  \nIn the case of more likely accretion from planetary embryos, additional silicate-sulfide separation could have occurred in magma oceans on both planets (Rubie et al. 2016), which were likely formed at later stages of accretion (see Elkins-Tanton 2012, Solomatov 2015, Schaefer and Elkins-Tanton 2018, Salvador et al. 2023, for reviews). The likelihood of magma ocean formation, their sizes, and evolution depend on the sizes and composition of impactors and timing. Formed impact-generated atmospheres impeded heat loss to space, allowed magma-atmosphere interfaces to exist, and extended the lifetime of magma oceans. Higher solar irradiation at lower heliocentric distances, as on Venus, resulted in elevated atmospheric and surface temperatures, thereby promoting the possible extended presence of a magma ocean (Hamano et al. 2013; Lebrun et al. 2013; Salvador et al. 2017). If Venus and Earth acquired equal amounts of water, the more substantial thermal blanketing effect of Venus' steam atmosphere could have played a significant role in the divergent evolution of the planets, affecting ocean solidification and degassing, mantle composition, and the fate of water. However, the formation of the steam atmosphere over magma oceans on Earth and Venus can be hindered by high H2O solubility in reduced ultramafic melts ( Sect. 16.4.2.1 ).  \nThe mass and fate of accreted water played a vital role in the physical-chemical evolution of magma oceans, affecting the behavior of elements during crystallization and degassing. It is possible that different amounts of water contributed to the divergent evolution of magma oceans on Earth and Venus (Salvador et al. 2023). Water affects the composition of outgassed volatiles (Gaillard et al. 2022), the melt-solid partition of elements, the mineralogy of forming mantle, ocean dynamics, and planetary tectonic style established after the crystallization (Schaefer and Elkins-Tanton 2018). Importantly, water could play a role in the oxidation state of magma and formed mantle.  \nWater was present in hydrated and oxidized carbonaceous chondritic materials (e.g., CI/CM type chondrites), and impact-generated atmospheres of terrestrial planets are modeled to be presented by H2, CO, H2O, CH4, NH3, H2S, and N2 (e.g., Schaefer and Fegley 2010; Zahnle et al. 2010, 2020). The modeled atmospheric compositions depended on the assemblage of impactors, magma-gas interactions, escape, and other factors (Schaefer and Fegley 2017). Late accretion of oxidized planetesimals could have increased mantle f O2, but the oxidation of the whole mantle on Earth cannot be explained. Photo-dissociation of atmospheric water, hydrogen escape, and net accumulation of remaining oxygen could have caused the oxidation of atmospheric gases (CO to CO2, H2S to SO2, etc.) and magma ocean tops, possibly affecting the entire magma oceans in a case of vigorous convection  \n<!-- image -->  \n(Hamano et al. 2013; Wordsworth 2016; Wordsworth et al. 2018; Krissansen-Totton et al. 2021). Regardless of the occurrence of the magma ocean, hydrogen escape could have been a valuable oxidation pathway of coupled atmosphere-crustal-mantle systems on Earth and Venus (Kasting et al. 1993; Sharp et al. 2013; Gillmann et al. 2009).  Zahnle and Kasting (2023) found that oxidation caused by H escape could be measurable to the upper mantle and crust if the upper mantle existed as a stagnant lid or if the magma ocean was stably stratified (not convecting). Despite the common assertion on an effective atmosphere-magma ocean exchange, it could be limited because of slower diffusion and higher viscosity in the lower temperature near-surface ocean and because of the likely occurrence of scum or froth material behaving as a stagnant lid.  In addition to the mass of affected magma via magma-gas partition, the degree of oxidation depends on the amount of accreted water, oxidation of iron species by water, and H2 production in the upper ocean via  \nFe 0 (metal) + H2O(g, in melt) → FeO (in silicates) + H2(g) (16.86)  \n3FeO (in silicates, melt, etc.) + H2O(g, in melt) → Fe3O4 (magnetite) + H2(g) (16.87)  \nAnother key factor is solar irradiation, which affects the greenhouse temperature, photodissociation of water, and escape of H and O. Formation of magmatic sulfate was only possible via oxidation by accumulated net atmospheric oxygen at the magma-atmosphere interface if the magma is not sufficiently mixed.  \nIrrespective of what happened to the water, high-pressure disproportionation of Fe 2+ to Fe 0 -metal and Fe 3+ (Hirschmann 2012; Schafer and Elkins-Tanton 2018; Armstong et al. 2019; Deng et al. 2020) in the lower magma oceans and sequestration of metal to the core could allow oxidation of silicate mantles on Earth-type planets. The disproportionation of Fe 2+ in the melt appears to be a more suitable process than the melting and recrystallization of bridgmanite advocated by Frost et al. (2004) and Wade and Wood (2005). The oxidation via Fe disproportionation does not go beyond the QFM buffer (Deng et al. 2020) and could have led to similar oxidation states in the mantles of Earth and Venus. The nearconstant f O2 of mantle rocks over the entire history of the Earth is consistent with this notion (Schafer and Elkins-Tanton 2018). The MnO/FeO ratio of Venus' surface samples (Table 16.3) does not suggest oxidation beyond the f O2 conditions of the QFM buffer (Schaefer and Fegley 2017; Sect. 16.2.2 ),  \n<!-- image -->  \nimplying S(-II) speciation in the magma ocean and sulfide mineralogy (pyrrhotite, troilite) in the mantle formed upon crystallization.  \nOn Venus and Earth, cores likely formed and compositionally evolved during planetary accretion. Jacobson et al. (2017b) modeled a multi-stage core formation on Earth-like planets, predicted the accumulation of O and Si in outer cores, and suggested that Venus could have better preserved such stratification due to the lack of a late giant impact that supposedly homogenized Earth's core ~100 Ma after CAIs (~4.47 Ga). They hypothesized that the lack of a detectable internally driven magnetic field on Venus reflects the chemical stratification of the core that prevented convection.  \nThe supposed planet-forming materials contain at least 0.5 wt% S as follows: ~2% in ordinary chondrites, ~2 -6 % in enstatite and carbonaceous chondrites, and ~0.5 -1% in iron meteorites (Jarosewich 1990; Goldstein et al. 2009; Lodders and Fegley 1998). The bulk S content of Earth is estimated to be 0.64 wt % (McDonough 2014), 0.9 wt% (Kargel and Lewis 1993), 2.9 wt% (Morgan and Anders 1980), and 0.61 ± 0.15 wt% (Wang et al. 2018, concordance estimate). However, the assessed S content in the bulk silicate Earth (BSE) is 200 -270 ppmw (e.g., O'Neill 1991; Kargel and Lewis 1993; McDonough and Sun 1995; Palme and O'Neill 2014; Wang et al. 2018; Kiseeva et al. 2024).  Sampled mantle and mantle -sourced igneous rocks suggest 120 -250 ppmw S in different mantle reservoirs (Kiseeva et al. 2024, for review). The discrepancy between S content in the Earth-building materials, bulk silicate Earth, and the mantle implies the segregation of sulfide melt to the core during a global silicate-Fe(metal) differentiation that likely occurred at the magma ocean stage. The depletion of chalcophile and highly siderophile elements (HSE) in mantle-derived rocks compared to abundances in chondrites is generally consistent with sulfide-silicate and metal-silicate partitioned during core formation (e.g., Wood et al. 2014; Kiseeva and Wood 2015).  \nA through segregation to the core remains to be reconciled with the chondritic-like S/Se/Te ratios and elevated S, Se, Te, and (HSE) abundances in mantle rocks. The standard explanation of mantle samples' S, Se, Te, and HSE abundances is an accretion of a veneer after core-mantle differentiation (e.g., RoseWeston et al. 2009; Wang and Becker 2013). Varas-Reus et al. (2019) demonstrate that the Se isotopic composition of mantle peridotites is consistent with a thin veneer of CI-type chondritic material delivered from the outer solar system. Tellurium stable isotope composition of mantle-derived rocks does not contradict a veneer of CM and/or enstatite chondritic material (Hellmann et al. 2021).  \n<!-- image -->  \nAlexander (2022) modeled that a thin (~0.2 wt% BSE) veneer of CM- or EL-like chondritic material could explain the abundance of elements but not the isotopic composition of S and Se. Wang et al. (2023) modeled that the chalcogen isotopic ratios in the Bulk Silicate Earth (BSE) could be influenced by evaporation from planetesimals. They concluded that the late veneer cannot exceed 0.2 wt% of the mantle's mass. Despite its small size, even a tiny veneer could be a significant nitrogen source in the BSE. An analogous late veneer is likely present on Venus. The lack of mantle homogenization by a late giant impact does not rule out a more significant contribution of veneer materials to the upper mantle composition and bulk nitrogen content. The abundant nitrogen in Venus' atmosphere (Table 16.12) supports this idea. To determine whether a later veneer of chalcogen-, highly siderophile element (HSE), and Nrich materials accreted, chemical and isotopic analysis of Venus' mantle -derived rocks and atmospheric chalcogens is required ( Sect. 16.5 ).  \nAlthough the segregation at the magma ocean stage likely involved separating immiscible FeS liquid from the silicate melt, different pathways regarding silicate-sulfide partitioning in magma ocean were modeled (e.g., O'Neill 1991; Rubie et al. 2016; Lauren z et al. 2016; Suer et al. 2021; Steenstra et al. 2022). The amount of precipitated sulfides depended on the magma ocean's sulfur abundance and solubility at sulfide saturation (SCSS), a function of temperature, pressure, melt composition (especially FeO content), and oxidation state (e.g., Baker and Moretti 2011; Smythe et al. 2017; Simon and Wilke 2024). Experiments below 24 GPa indicate decreasing SCSS with increasing pressure and decreasing temperature (e.g., Blanchard et al. 2021). The low S solubility at the depth of a magma ocean supposedly allowed the exsolution of droplets of liquid FeS (socalled 'the late Hadean matte') and their segregation to the core (O'Neill 1991). The Earth's magma ocean crystallization model of ZhanZhou et al. (2024) based on generalized SCSS experimental data at <24 GPA (< 746 -970 ppmw) also suggests latestage sulfide precipitation from evolved (FeO-SiO2-Al2O3 enriched) magma at depths of 120 -220 km that does not suggest sulfide percolation through the crystallized mantle.  Overall, the percolation of liquid sulfides to the core is the standard explanation for the abundance of chalcophile and siderophile elements in mantle-derived rocks (Kiseeva et al. 2024, for review).  \nData on sulfur solubility at deep magma ocean pressures (43 -53 GPa) from Steenstra et al. (2022) revealed a high SCSS (0.3 -0.4 wt% S). Their model suggests a consistent SCSS throughout the ocean, allowing sulfide segregation only at the final stages of magma crystallization, and implies the formation of a relatively S-rich mantle. In this case, the abundance of siderophile elements in mantle rocks may be  \n<!-- image -->  \ncaused by Fe disproportionation rather than sulfide segregation to the core. If the contentious model of Steenstra et al. (2022) is correct, sulfide-rich regions in the upper parts of the mantle with FeS trapped in the silicate matrix could have formed during the late stages of ocean crystallization. The slow solidification of Venus' magma ocean (~100 Ma compared to Earth's ~5 Ma, Hamano et al. 2013, 2024) could have affected the formation of a sulfide-enriched mantle and possibly enhanced the development of exceptionally S-rich upper mantle regions.  \nOn Earth and Venus, the ocean stage and subsequent occurrence of mantle sulfides were influenced by the accretion of various S-rich impactors, differentiated (core-bearing) and chondritic. These impactors enhanced sulfide segregation and established the abundances of sulfur and chalcophile and siderophile elements in the mantle (Rubie et al. 2016; Suer et al. 2017; Li et al. 2016). Diverse pathways are modeled for multi-stage (Rubie et al. 2016) and homogeneous (Suer et al. 2017) accretion. The potential ab sence of a giant impact homogenization of Venus' mantle may have preserved an S -rich mantle and sulfide-rich upper mantle regions formed via magma ocean crystallization. This hypothesis could be tested through in situ measurements of S and chalcophile and siderophile elements in unaltered basalts ( Sect. 16.5.2 ). Potential sulfide-rich upper mantle regions could be inferred from gravity data and by investigating large impact crater ejecta ( Sect. 16.5.2 ).  \nThe different mantle convection patterns on Venus (Rolf et al. 2022) and Earth could have affected further segregation of sulfides to cores. Sulfides formed via later magma ocean crystallization could remain in the lithospheric lid. If basal magma ocean ex isted throughout Venus' history (O'Rourke et al. 2020), it could be enriched in dissolved sulfur at the SCSS level and contain FeS liquid accumulated at the bottom, as in some mafic magma chambers on Earth ( Sect. 16.2.2 ). If Venus did not have a large-scale magma ocean, melting silicates in interiors due to radioactive heating allowed the separation of sulfidemetal liquids and their gravitational accumulation to cores. Such a pathway, but with 26 Al decay as an energy source, likely formed metal-sulfide cores on many differentiated planetesimals without the formation of magma oceans (e.g., McCoy and Bullock 2016; Elkins-Tanton 2016).  \nThe composition and oxidation states of atmospheres formed via magma ocean degassing have been the subject of interest.  If magma ocean convection is limited, the outgassing of oxidized volatiles (CO2, H2O) could have resulted from oxidation of the ocean tops caused by hydrogen escape. If the deep ocean was oxidized via disproportionation of Fe(II), reduced magmatic gases (CO, H2) could have been  \n<!-- image -->  \nreleased into the atmosphere. However, a vigorous mixing of an ocean equilibrated with Fe-metal at depth could have created a positive f O2 gradient below ~50 -200 km, allowing a release of oxidized magmatic gases (CO2, H2O) at f O2 around the QFM buffer (Armstrong et al. 2019; Deng et al. 2020), and SO2 could be among the gases as well.  \nChemical co-evolution of the atmospheric-ocean system could have led to more oxidized atmospheric compositions at the late stages of accretion. Hirschmann (2012) suggested a gradual depletion of atmospheric CO2 via dissolution in more oxidized near-surface magma followed by precipitation of graphite or diamond in the interior of the convecting magma ocean. The higher abundance of degassed CO2 on Venus (Table 16.12) could indicate less such trapping than on Earth (Armstrong et al. 2019), possibly due to the lack of a large magma ocean created by violent impact(s). By analogy with CO2, an excessive atmospheric SO2 could be dissolved at the atmosphere-magma interface and reduced to sulfides at depth.  \nGaillard et al. (2022) included SO2 and H2S in outgassing models for shallow magma oceans. At oxidizing conditions (>IW + 2, i.e. >2 log f O2 units above the IW buffer), they predicted the exsolution of CO2, SO2, N2, and H2O to a dense ~100-bar atmosphere. S-poor and H2-CO dominant gases were modeled for reduced magmas (<IW - 2). Even moderately reduced melts (<IW + 1) release only trace H2S (< 0.05 vol%) due to the high solubility of sulfur.  Gaillard et al. (2022) illustrated that the Earth's crustal inventory of sulfur (1.2 × 10 22 kg, Canfield 2004) could be justified by degassing from oxidized (~IW + 2.5) magma ocean. In application to Venus, this model suggests that the apparent accumulation of sulfur in atmospheric and crustal materials ( Sect. 16.2 ) required relatively oxidizing (>IW + 2) conditions in the upper magma ocean and/or in the history of volcanic degassing ( Sec t. 16.4.3 ). The occurrence of supposedly volcanic and abundant sulfur species in the present atmosphere-surface system aligns with relatively oxidized magma and parent mantle rocks.", '16.4.2  Aqueous and Anhydrous Pathways': "The potential for water condensation and the destiny of a possible hydrosphere on early Venus have been evaluated (Ingersoll 1969; Kasting and Pollack 1983; Kasting 1988; Abe and Matsui 1988; Way et al. 2016; Salvador et al. 2017; Way and Del Genio 2020; Krissansen-Totton et al. 2021). Sulfur has not been specifically discussed in Venus' aqueous conditions, and the following inquiries remain to be addressed.  \nWas volcanic-sourced SO2 oxidized to H2SO4 in the moist atmosphere and disproportionate to sulfate and native sulfur in aqueous environments at the planetary surface? What were the sulfur isotopic signatures of such processes? Was trapping of volcanic-sourced SO2 and minerals sufficient to prevent the formation of sulfuric acid clouds? Did the atmosphere become oxidized sufficiently to allow for sulfate accumulation in oceanic water and the formation of sulfate-bearing sediments? Are pyrite and/or native sulfur formed via aqueous oxidation of magmatic sulfides and sulfate reduction?  Was native sulfur reduced, oxidized, and disproportionate to sulfides and sulfates?  \nIf water avoided condensation, the inquiries are related to sulfur in a hot, evolving atmosphere-crustal system with varying water content. These inquiries include the fate of sulfuric acid clouds, the oxidation of the atmosphere and crustal materials driven by hydrogen escape, and the partitioning of volcanically degassed sulfur between the atmosphere and solid materials, along with the corresponding isotopic fractionation. These sulfur-related problems have not been addressed in the literature. This section outlines the likelihood of liquid water being present on Venus and first discusses the implications for sulfur on the presumably abiotic planet.", '16.4.2.1   Did Water Condense?': "The possibility of similar initial water abundances on Earth and Venus ( Sect. 16.4.1 ), the presence of moderately oxidized igneous rocks at the landing sites of the Venera and Vega spacecraft ( Sect. 16.2.2.1 ), along with the high atmospheric D/H ratio, indirectly point to a period in Venus' history when water was abundant in outer shells of the planet. However, no data indicates the presence of past liquid water, and there is no consensus on this topic.  \nVenus could have accreted less water than Earth, which constrained water condensation if water was released into the atmosphere (e.g., Kasting 1988). The planet could have been desiccated through Fe 0 -steam reactions during a collision of large embryos that also changed Venus' spin, as Davies (2008) hypothesized. Such a pathway implies H2 formation via (16.86) and (16.87) and loss before the solidification of the planet and oxidation of its interior. Gillmann et al. (2016) illustrated that postaccretion impacts of 400-800 km bodies could have triggered volcanism, heating, and atmospheric erosion, placing the planet in a runaway greenhouse. An early atmospheric H2O could be lost via photodissociation by intense solar UV flux from the young Sun and following hydrogen and, in some  \n<!-- image -->  \nconditions, oxygen escape (Zahnle and Kasting 1986; Hunten 1993; Chassefiére 1997; Kulikov et al. 2006; Lammer et al. 2008, 2011; Gillmann et al. 2009; Lichtenegger et al. 2016; Zahnle and Kasting 2023). A portion of atmospheric steam could be lost through oxidation of Fe 0 -metal and silicate's Fe(II) in impactgenerated gases, steam-magma, and steam-silicate interactions (16.86)  and (16.87), followed by hydrogen escape. As an illustration, the interaction of dissolved H2O with Fe(II) in crystallizing seafloor basalts leads to the formation of magnetite and H2 (Holloway 2004).  \nIn one modeled scenario, Hamano et al. (2013) demonstrated that thermal blanketing of steam atmosphere and solar irradiation at ~0.7 AU allowed a slow magma-ocean solidification within ~100 Ma. They modeled that corresponding slow water outgassing could have desiccated the planet through hydrodynamic hydrogen escape. Their models showed that the mantle water reservoir becomes more extensive upon ocean solidification than the water reservoir in the atmosphere, limiting chances for condensation. Models of Majumder and Hirschmann (2017), Miyazaki and Korenaga (2022), Bower et al. (2022), Sossi et al. (2023), and Maurice et al. (2024) also suggested water holding due to inefficient magma ocean outgassing during crystallization. Gaillard et al. (2022) noted suppressed degassing of high-solubility magmatic water in a 100-bar CO2/CO-rich atmosphere. As described in Holloway's (2004) study on submarine lava, trapped magmatic water could be consumed via melt oxidation (Eq. (16.87)) at the uppermost layer of the magma ocean, contributing to the oxidation of the mantle. This is a plausible and unaddressed mechanism of water consumption in the upper magma ocean on Venus and elsewhere.  \nContrary to Earth, possible fewer large-scale collisions neither contributed to the formation of an extensive magma ocean nor aided in water degassing ( Sect. 16.4.1) . After the crystallization of Venus' magma ocean, the possible establishment of the stagnant lid mantle convection pattern (Rolf et al. 2022) could have limited the magmatic degassing of water. The limited degassing could be compensated by hydrogen escape without water condensation (Miyazaki and Korenaga 2022). The preservation of water in the mantle is consistent with the petrological interpretation of Vega 2 rock composition (Table 16.3) (Barsukov et al. 1986b; Barsukov 1992). A few oceanic masses could be stored within Earth's nominally anhydrous mantle phases (e.g., Kaminsky 2018), and Venus' mantle might be similar (Zolotov et al. 1997).  \nThe high atmospheric D/H ratio can be explained without invoking liquid water at the surface (Greenspoon and Lewis 1988; Donahue et al. 1997; Zahnle and Kasting 2023). Oxidation of Venus' upper interior could have occurred without hydrogen escaping from the steam atmosphere ( Sect. 16.4.1 ), and moderately oxidized igneous rocks at landing sites ( Sect. 16.2.2.1 ) may not indicate past water-rich environments.  Venus' one-modal hypsometric curve does not indicate the existence of large cratons ( Sect. 16.2.2.2 ) that could have formed in the presence of water ocean (Campbell and Taylor 1983). The supposedly layered rock formations in tessera terrains ( Sect. 16.2.2.4 , Figs. 16.12, 16.13) might be of volcanic origin rather than sedimentary rocks deposited by water.", '16.4.2.2  Atmospheric and Crustal Sulfur Without Water Condensation': "The fate of sulfur in the runaway greenhouse atmosphere and upper crustal rocks depended on the amount of atmospheric water, solar UV irradiation, volcanic degassing, and resurfacing by volcanism and tectonics.  The amount of atmospheric water was determined by accretion, early evolution ( Sect. 16.5.1 ), ongoing volcanic degassing, and consumption via hydrogen escape and steam-rock reactions, causing oxidation and hydration (e.g., formation of amphiboles). The amount of atmospheric water and cloud coverage mainly determined the atmospheric and surface temperature. The amount of atmospheric water and solar UV flux affected the photodissociation of water vapor, fluxes of H and O to space, and the net mass of O available for oxidation of atmospheric and crustal materials (e.g., Zahnle and Kasting 2023). Regardless of the redox state of the lower atmosphere and surface, oxidation conditions in the upper atmosphere due to CO2 and H2O photolysis favored the oxidation of S-bearing gases to S(VI) species (SO3, H2SO4(g), sulfuric acid aerosol). SO2 dominates over reduced S-bearing gases (Table 16.1), and the likely case at higher water abundance allowed more oxidizing conditions via H escape. Correspondingly, sulfatization rather than sulfidation of surface materials provided the sink of atmospheric sulfur from a water-rich atmosphere.  \nBecause the amount of Ca and Na in exposed volcanic rocks exceeds the S gases degassed from them ( Sect. 16.4.3 ), most of the degassed sulfur must have been trapped in secondary minerals, and a complete sulfurization of permeable volcanic rocks was unlikely because of that disbalance. Any increase in atmospheric sulfur mass due to volcanic degassing was followed by consumption by gas-solid reactions, and prolonged periods without volcanism led to lower f SO2 that approached values controlled by equilibria with sulfates (e.g., (16.75, 16.85)). As in the present epoch, the presence of even trace  \n<!-- image -->  \nvolcanic SO2 in the atmosphere likely prevented the formation of carbonates via weathering reactions. If low temperature and f SO2 allowed stability of carbonates, their formation would be hindered by slow CO2-solid reactions. In weathering crusts, carbonates could not have survived episodes of large-scale volcanic degassing of S-bearing compounds ( Sect. 16.4.3 ) and oxidation of sulfides in lava that enhanced sulfur abundance in the atmosphere.  \nIn the case of an H2O-poor atmosphere, the greenhouse temperature and the fate of sulfur species in the atmosphere-crust system might not have been much different from that in the current epoch that began after the last global volcanic resurfacing ( Sect. 16.4.4 ). The disproportionation of sulfur in SO2solid reactions (Eq. (16.62)) likely led to the formation of metal sulfates and S2 gas.  Pyrite could have formed via pyrrhotite oxidation and sulfidation of Fe(II) in exposed glasses and silicates ( Sect. 16.3.2.3 ). Even if S-MIF occurred in O2-less upper atmosphere, the corresponding signature might not have been recorded in rocks due to the evaporation of condensed Sx in the lower atmosphere.  \nBoth elevated H2O(g) contents and moderate UV fluxes (Zahnle and Kasting 2023) enhanced the escape and H/O ratio in escaping gases. These factors allowed elevated S(VI)/(S(IV) + S(0) + S(-II)) and CO2/CO ratios in atmospheric gases and favored accumulation of atmospheric O2. SO2 was more than SO3 except for the extreme O2-rich (> 1 bar) and hot (>1000 K) conditions (Fig. 16.26). If oxygen escape was limited, the evolution of a steam atmosphere could have led to the oxidation of a significant Fe(II) mass in rocks (Zahnle and Kasting 2023), and the supposed atmospheric O2 should have caused geologically instantaneous oxidation of magmatic sulfides (pyrrhotite, pentlandite, etc.) in exposed materials,  \n2Fe0.875S + 3.3125O2 → 0.857Fe2O3 + 2SO2 (16.88)  \nAccording to experimental data (e.g., Alksnis et al. 2018; Aracena and Jerez 2021; Liu et al. 2023), the oxidation of ferrous sulfides to hematite is favored by elevated temperature and occurs via the formation of magnetite as an intermediate phase. Although the released SO2 contributed to the atmospheric inventory of sulfur and favored the formation of sulfuric acid clouds, the inventory was controlled by a balance between sources and sinks to minerals. The sinks were enhanced by elevated surface temperature and occurred by exemplary net reaction,  \nSO2 + 0.5O2 + CaSiO3 (in pyroxenes and glasses) → CaSO4 + SiO2 (16.89)  \nin which SO2 and O2 could be replaced by SO3. In contrast to present anoxic surface environments ( Sect. 16.3.2 ), sulfatization occurred without sulfur disproportionation, and sulfur did not return to the atmosphere. The crustal thickness affected by anhydrous sulfatization and the degree of sulfatization depended on volcanic and tectonic resurfacing, though aqueous environments ( Sect. 16.4.2.3 ) allowed more efficient resurfacing via erosion and sedimentation.  \nHydrogen scape and a limited O escape in the atmosphere with one Earth's ocean mass likely caused a large-scale oxidation of crustal materials (e.g., Kasting 1988; Zahnle and Kasting 2023). A significant greenhouse in the steam-rich atmosphere favored crustal melting (Kasting 1988), and the increased geothermal gradient facilitated upper mantle melting and corresponding volcanic activity. Although a widespread occurrence of surface magma allowed trapping of net atmospheric O through oxidation CO2, H2O, and O2, atmospheric f O2 could not have strongly exceeded f O2 values controlled by the Hem-Mt buffer (16.66), that is, 10 -5 to 10 -3 bars. The plausible buffering of atmospheric redox state by the MtHem equilibrium in surface magmas set the atmospheric SO3/SO2 ratio of ~10 -3 (Fig. 16.26), CO2/CO ratio, and allowed sulfur oxidation in affected silicate melts. Magmatic sulfur is more soluble in S(VI) form (e.g., Simon and Wilke 2024), and oxidation of the surface melts in contact with a hot steam atmosphere enhanced sulfur solubility. Magmatic S(-II) was oxidized to S(VI), and atmospheric sulfur (SO2, SO3) was trapped in the sulfate melt complexes.  The enhanced solubility also limited the degassing of magmatic sulfur supplied to magma's molted surface ponds.  Together with high-temperature sulfurization of the surface, such trapping of volcanogenic sulfur oxides confined the formation of sulfuric acid clouds. Solidification of surface melts led to anhydrite-bearing igneous rocks. In addition to abundant ferric oxides (Zahnle and Kasting 2023), anhydrite is expected to be present in rocks that consume the O mass as in the Earth's ocean, and no S-MIF signature is anticipated. Such rocks' direction would indicate large-scale oxidation facilitated by the steam greenhouse environments.", '16.4.2.3   Water Condensation and Aqueous Environments': "Feasible Earth-like initial water inventory, an accumulation of water in a primordial atmosphere through impact and magma degassing ( Sect. 16.4.1 ), and the low luminosity of the young Sun favored condensation of atmospheric water after accretion and cooling of the surface.  Although water vapor and albedo played crucial roles in determining the surface temperature and condensation, limited CO2 degassing or its sequestration as graphite (Hirschmann 2012) marginally mitigated greenhouse heating.  \nRadiative-convective atmospheric modeling (Kasting and Pollack 1983; Kasting et al. 1984; Kasting 1988; Abe and Matsui 1988) showed that the early solar flux at the Venus' orbit was close to the critical value required to trigger the runaway greenhouse (Ingersoll 1969; Rasool and de Bergh 1970). Abe et al. (2011) and Salvador et al. (2017) showed that an albedo slightly higher than that on today's Earth is needed to have an ocean on early Venus. In models of Kasting (1988), Way et al. (2016), Way and Del Genio (2020), and Salvador et al. (2017), high-albedo clouds cooled the surface and favored condensation. Models show that clouds and the faint early Sun could have allowed a long (~0.5-3.8 Ga) presence of liquid water (Pollack 1971; Kasting 1988; Abe and Matsui 1988). These were 1-D models that had to make assumptions about the albedo, which were unconstrained. Yang et al. (2014), Way et al. (2016), and Way and Del Genio (2020) emphasized the importance of the slow rotation of Venus, rather than solar luminosity or water mass, to allow surface water over extended periods of its history. The speculated felsic composition and layered rocks of Tessera terrains ( Sect. 16.2.2.4 ) and possible chloride melts that formed canali ( Sect. 16.2.2.3 ) indirectly suggest the formation of water-deposited sediments and evaporites.  \nIf atmospheric water condensed (Pollack 1971; Kasting 1988; Abe and Matsui 1988; Yang et al. 2014; Way et al. 2016; Way and Del Genio 2020; Salvador et al. 2017; Krissansen-Totton et al. 2021), Venus' surface temperature was below the critical temperature of water (647 K or less for salty water). The cloud-free moist greenhouse models suggest temperatures below ~500 K, while models that considered elevated albedo of clouds resulted in temperatures as low as ~300 K (Pollack 1971; Kasting 1988). Global three-dimensional simulations considering clouds, slow rotation, and topography suggest ~276-420 K, depending on location and elevation (Way et al. 2016; Way and Del Genio 2020). Therefore, ambient conditions can span from sub-critical hydrothermal to temperate environments typical for most of Earth's history.  \nIn the moist greenhouse, atmospheric warming is limited by growing planetary albedo at increasing solar luminosity. Together with diffusion-limited hydrogen escape, this extends liquid water's existence (Kasting 1988). In his model, hydrogen escape and net oxygen accumulation become efficient above ~150 K. Kasting (1988) and Abe et al. (2011) noted that a significant fraction of ocean water could be lost via hydrogen escape before the true runaway occurs. Oxygen left after hydrogen escape was supposedly consumed via the oxidation of C-, Fe-, and S-bearing species in the atmosphere, oceanic water, and the crust.  \nAs on Earth, aqueous processes on land, oceanic, and hydrothermal settings should have included erosion, transport, and deposition of solids, chemical alteration (dissolution, hydration, oxidation, etc.) of minerals and glasses, transport of ions, and precipitation of secondary minerals in situ and from transferred solutions. The surface and oceanic water compositions were formed by dissolving minerals and glasses that supplied major cations (Na + , Ca 2+ , Fe 2+ , etc.) and volcanic/impact degassing that provided anions (Cl -, HS -, HCO3 -, etc.).  \nIn the case of an initial anoxic atmosphere, NaCl-type ocean water with dissolved inorganic carbon (DIC, CO2 + HCO3 -+ CO3 2) was likely. The concentration and speciation of DIC depended on atmospheric f CO2, temperature, and pH , which reflected acid-base balances in the water-rock systems. The oceanic pH was likely neutral to alkaline, as carbonic acid forming via CO2 dissolution would be neutralized by cations, leading to carbonate precipitation at neutral or alkaline pH . A CO2-rich atmosphere could not sustain in the long run (~10 7 -10 8 a) when in contact with mafic/ultramafic rocks, mainly if new rocks were introduced and carbonates were removed from the surface through subduction or burying (e.g., Walker 1977; Sleep and Zahnle 2001; Zahnle et al. 2007; Zolotov 2020).", '16.4.2.4   Sulfur in Aqueous Environments': "Sulfur behavior on early Venus may differ from that on early Earth due to the absence of life (e.g., sulfate-reducing, sulfide-oxidizing, Fe 2+ -oxidizing, and O2-producing species) and biological organic matter. However, abiotic organic-poor high-temperature settings may not differ significantly from Earth's counterparts. As on Earth, the solubility of iron sulfides likely controlled the concentration of major S-bearing solutes (H2S and HS -, depending on the pH and temperature) in surface and oceanic water. The solubility increases with temperature (Pokrovsky 2024), and sulfides in igneous rocks (pyrrhotite, pentlandite) could be strongly affected by dissolution in ambient high-temperature (> ~500 K) environments and hydrothermal systems. Higher greenhouse temperature and lower pH at elevated atmospheric f CO2 enhanced concentrations of dissolved H2S in surface and ocean-top water.  Secondary sulfides of iron (pyrite, marcasite, pyrrhotite) and trace metals (Cu, Zn, Pb, Co, etc.) could have precipitated in ambient high-temperature settings on land and Earth-like basalt-hosted hydrothermal systems throughout the globe. As in Earth's hydrothermal systems (Kump and Seyfried 2005), the supply of Fe 2+ via local high-temperature sulfide dissolution favored the precipitation of secondary sulfides in  \n<!-- image -->  \nlower temperatures and more alkaline settings. As on Earth, pyrite could be the most significant secondary sulfide precipitated at elevated bulk sulfur content and higher pH (Fig. 16.27).  \nThe aqueous oxidation of sulfides on Venus has not been evaluated, and only qualitative inferences could be made based on analogs and insights from equilibrium chemical thermodynamics and kinetics of abiotic redox reactions. In general, the sulfate concentration in surface and oceanic water could have been influenced by a dynamic balance involving volcanic/impact degassing of S-bearing species, their photochemical and chemical oxidation, sulfide oxidation, and trapping through precipitation and hightemperature abiotic reduction.  \nOn Earth, sulfate content in the Archaean seawater did not exceed 2.5 micromoles (Crowe et al. 2014) despite some initial oxidation of sulfides on land.  During the Great Oxidation Event (GOE, Holland 2006) at ~2.4 Ga, atmospheric x O2 abruptly reached approximately 10 -2 after nearly 0.5 Ga of oxygenic photosynthesis in late Archaean (Lyons et al. 2014). The large-scale oxidation of sulfides on land occurred during the event and enhanced oceanic sulfate content. Whether free O2 accumulated on Venus depended on relative rates of hydrogen escape and net oxygen sinks via oxidation of atmospheric and volcanic ( Sect. 16.4.3 ) gases (CO, SO2, OCS, H2S, Sn, and H2), aqueous (Fe 2+ , H2S, HS -), and crustal ferrous materials. The relative masses of escaped hydrogen, Fe(II), and sulfides available for oxidation played crucial roles in determining the fates of oxygen, iron, and sulfur in the coupled atmosphereocean-crust systems.  \nSome sulfates could have formed regardless of the O2 accumulation in the lower atmosphere, which is possible through a rapid hydrogen escape from a water-rich atmosphere-ocean system. As on early Earth (Farquhar et al. 2001; Ono 2017) and Mars (Frantz et al. 2014; Frantz 2024), photolysis of SO2, OCS, or H2S in anoxic atmosphere could have caused the formation of native sulfur and S(VI) species (SO3, SO4 2) leading to a S-MIF isotopic signature in surface waters and corresponding sediments. In a water-saturated atmosphere, the photochemical oxidation of SO₂ (Eqs. (16.6), (16.7)) and other Sbearing gases ( Sect. 16.3 ) must have produced sulfuric acid aerosols and acid rain. However, this process could be limited if the sink of atmospheric sulfur exceeds the supply of reduced compounds (SO2, OCS, Sn, H2S). An efficient volcanic SO2 ( Sect. 16.4.3 ) and SO2-bearing impact plumes generated by hydrated carbonaceous chondrite impactors (Schaefer and Fegley 2010) and/or in water-bearing targets followed by acid rains could have caused acid sulfate weathering on land, somewhat like early Mars (e.g., Franz  \n2024; Zolotov and Mironenko 2016). Dissolution of SO2 in water caused the formation of sulfurous acid (H2SO3) and bisulfite ions (HSO3-) in acid rain and surface water. Disproportionation of S(IV) in latter compounds led to SO4 2and S(0) and/or H2S depending on  temperature, pressure, pH , and catalytic conditions  \nSO2(aq) + H2O(l) → H2SO3(aq) → SO4 2+ 2H + + H2S(aq) (16.90)  \nAs on Earth, S(IV) disproportionation could have occurred in ambient, diagenetic, and hydrothermal environments. In contrast to diverse bio-mediated processes (Fisher 2008), abiotic low-temperature disproportionation was less efficient.  \nSulfates of Ca, Mg, Na, and K precipitated in the weathering crusts together with clay minerals, migrated in surface and ground waters, deposited via evaporation (c.f., Zolotov and Mironenko 2016), and delivered SO4 2ions to oceans. Kinetically inhibited aqueous sulfate reduction below ~500 K (Ohmoto and Lasaga 1982) allowed metastable coexistence of SO4 2with H2S and HS -provided via the dissolution of sulfides and volcanic H2S. Although sulfates could be more abundant at ocean tops, opposite vertical gradients of sulfate and sulfides may not be as strong as in the Black Sea, which is rich in biological organic compounds in subsurface waters and bottom sediments.  \nExcept in alkaline conditions, oxidation of aqueous sulfides to SO4 2requires lower f O2 than Fe 2+ oxidation to low-solubility ferric oxyhydroxides (Garrels and Chryst 1965), Fig. 16.27. On Venus, sulfide oxidation in weathering crusts affected by lowpH rains was possible before a significant increase in atmospheric f O2 and thorough oxidation of oceanic Fe 2+ . Even low-temperature oxidation of aqueous sulfides by O2 occurs rapidly (Millero et al. 1987). Therefore, an increase in Venus' atmospheric f O2 could have caused the oxidation of sulfides to SO4 2at the ocean tops and on the land following the transfer of sulfate- and Fe 2+ -bearing waters to oceans, where they coexisted (Fig. 16.27) until further f O2 growth.  \nDuring the GOE, the influx of sulfates from continents formed via the oxidation of sulfides facilitated the microbiological reduction of oceanic sulfate and enhanced sulfide content in the deeper waters. It boosted pyrite formation in organic-rich bottom environments (Reinhard et al. 2009). The biological reduction of sulfate could have restricted the accumulation of oceanic sulfate, which had remained unabundant during Archaean (Habicht et al. 2002; Crowe et al. 2014). Despite the supply of sulfates to the ocean, deep oceans partially remained reduced (Fe 2+ -rich, sulfidic, and sulfate-poor) until after ~1.8  \nGa (Isley 1995; Kump and Seyfried 2005). Opposite to Earth's marine sediments (Rickard 2012b; Emmings et al. 2022), the apparent lack of biological sulfate reduction and abundant organic matter on Venus favored the accumulation of oceanic sulfates and could have extended the lifetimes of SO4 2and Fe 2+ in sulfate-poor water before a significant O2 buildup.  \nOn both planets, oceanic sulfate could be withdrawn through reduction above ~470-520 K in hydrothermal systems hosted in mafic and ultramafic rocks (Alt et al. 1989) and precipitation of pyrite  \nSO4 2+ 2H + + 4H2(aq) → H2S(aq) + 4H2O(l) (16.91)  \nFe 2+ + 2H2S(aq) → FeS2 + H2 + 2H + (16.92)  \nMore reduced venusian rocks ( Sect. 16.2.2.1 ) favored sulfate reduction by providing more H2 (Eq. 16.87) and Fe 2+ for reaction (16.92). If Venus had higher ambient and hydrothermal system temperatures, they favored abiotic sulfate reduction (16.91).  \nAnother aqueous sulfate sink is a high-temperature anhydrite formation in basalt-hosted oceanic hydrothermal systems where Ca 2+ is sourced from basalts (Alt et al. 1985, 1989; Chen et al. 2013), also inferred in early Venus' models (Zolotov and Mironenko 2009). The high temperature favored anhydrite deposition in surface and hydrothermal settings and precipitation of high-solubility Na and Mg sulfates via evaporation in lagoons. The latter process provided the major sink of sulfates upon exhaustion of surface water.  \nAs on Earth (Walters et al. 2020; Schwarzenbach and Evans 2024), the subduction of oceanic lithospheric plates could have delivered sulfur to the mantle, affecting its redox state and sulfur content. However, the unknown likelihood of plate tectonics on Venus hinders implications. Despite sulfate formation in photochemical and lowf O2 aqueous settings, the concentration of oceanic sulfates could be comparable to that in Archaean oceans (Habicht et al. 2002; Crowe et al. 2014). As in Earth's history, developing sulfate-rich seawater required significant oxidation of land sulfides and oceanic HS -and H2S. Such oxidation was questionable without a substantial accumulation of atmospheric O2.  \nThe GOE occurred due to an imbalance between net O2 sink via oxidation and production due to the burial of organic C and escape of H to space (Catling 2014; Lyons et al. 2014). Venus lacked oxygenic photosynthesis, which provides the primary net O2 supply on Earth, and several factors could have offset  \n<!-- image -->  \nthe supply of O2 from the upper atmosphere. The slow and prolonged (~100 Ma, Kasting 1988) net oxygen production in the moist greenhouse atmosphere due to limited hydrogen escape could have allowed rapid abiotic oxidation of aqueous Fe 2+ (Davison and Seed 1983) and sulfide species (Millero et al. 1987) on the land and at ocean tops. Without felsic continents ( Sect. 16.2.2 ), a prevalence of Fe(II)rich mafic and ultramafic rocks throughout the planet favored O2 consumption. Dominant mafic/ultramafic rocks on land and ocean floors provided more Fe(II) for oxidation than in Archaean. More reduced mafic Venus' rocks ( Sect. 16.2.2.1 ) provided more capacity to consume O2 than Earth's counterparts. On early Earth (Isley 1995; Kump and Seyfried 2005), seafloor hydrothermal supply of Fe 2+ strongly contributed to the ferrous iron oceanic inventory. An advanced hydrothermal supply of Fe 2+ and H2S is envisioned for Venus, providing higher temperatures and rock's Fe(II)/Fe(III) ratios.  \nWhether Venus' atmospheric O2 reached a percentage mixing ratio depended on the balance between the O2 flux from the upper atmosphere and the supply of reductants. A significant accumulation of O2 could have occurred at a limited supply of volcanic rocks and hydrothermal Fe 2+ and H2S. Limited resurfacing on land and the lack of plate tectonics that removed oxidized materials favored the accumulation. Other factors that could limit oxidation are oceanic stratification and the lack of Fe 2+ oxidizing microorganisms. Enhanced f O2 facilitated the precipitation of ferric oxyhydroxides in the ocean and the oxidation of sulfides on land and in the ocean. On Earth, episodic precipitation of ferric species through biological and abiotic oxidation of oceanic Fe 2+ led to banded iron formations (BIF) formed before, during, and after the GOE (Isley 1995; Yin et al. 2023). Correspondingly, Venus' BIF-like formations could have formed in a range of atmospheric f O2.  \nMass balance assessments show that trapping the oxygen remaining after H loss to space from one Earth's ocean mass of water requires Fe(II) oxidation to hematite in a global layer of several tens of km (Kasting and Pollack 1983; Lécuyer et al. 2000; Zahnle and Kasting 2023). Aqueous environments provided much more valuable conditions for consuming the vast mass of oxygen than gas-solid interactions. If Venus' atmosphere lost such an amount of water during the moist greenhouse, one would expect the development of abundant BIF-like formations via the oxidation of Fe 2+ -rich oceanic water and their significant burying in the crust and/or the mantle. The formation of sulfate-rich oceanic water would likely occur if a water mass comparable to Earth's ocean had been lost. Fe(II) oxidation within several tens of km implies a significant oxidation of sulfides in those rocks.  \n<!-- image -->", '6.4.2.5   The Cessation of Aqueous Period': "If water condensation ever occurred, the subsequent consumption of surface water would have influenced the fates of sulfur and other elements (Cl, Na, Fe, Si, etc.) involved in aqueous processes. Net consumption of liquid water likely occurred via hydrogen escape from the moist greenhouse atmosphere (Kasting 1988) and the hydration of crustal materials. Some water might have been removed by the subduction of hydrated slabs into the mantle, thereby preserving hydrogen in highpressure phases (Albarède 2009). These processes desiccated increasingly warming environments, facilitating evaporation, brine formation, and precipitation. Alteration of rocks by sub-critical oceanic water strongly affected the water and rock composition. The precipitation of low-solubility anhydrite further depleted oceanic sulfates. The enhanced dissolution of iron sulfides and silicates led to chloride-, silica-, and sulfide-rich oceanic water, as illustrated in initial models of Zolotov and Mironenko (2009). The evaporation of remaining water placed the planet into a runaway greenhouse state (e.g., Ingersoll 1969; Kasting 1988). The consumption of liquid water led to silica-rich rock and chloride evaporites forming on dry ocean floors (Zolotov and Mironenko 2009) that could be distinct from sulfate-rich facies deposited earlier at lower temperatures.  \nHow the aqueous period ended is still being determined. In one commonly considered pathway, the surface temperature approached the critical point for water as solar luminosity increased. In Way and Del Genio's (2020) model, increasing solar luminosity does not significantly affect ocean-bearing Venus's global mean surface temperature. They suggested massive volcanic degassing episode(s) that increased the greenhouse temperature and caused ocean evaporation. One can suggest that an impact-caused temperature spike could trigger a runaway greenhouse. Extreme runaway greenhouse heating (e.g., Kasting et al. 1988) favored decarbonization (e.g., Höning et al. 2021) and dehydration of crystal materials that contributed to warming.  \nSilicate rocks could have melted at the surface (Kasting 1988; Warren and Kite 2023) and in the upper mantle, triggering volcanism. Partial or complete melting of chlorides, silica phases, sulfides, silicates, sulfates, and iron oxides (e.g., from BIF-like rocks) favored the assimilation of compounds formed in the aqueous period.  If basalts melted, redox reactions led to melts with higher Fe(III)/Fe(II), S(VI)/S(-II), and S/Cl ratios and SiO2 contents than in initial basalts. Magma's Fe(II) could be oxidized by oxygen formed via hydrogen escaping from the atmospheric water (Warren and Kite 2023). Additional magmatic  \n<!-- image -->  \nsulfates could have formed if near-surface magma was oxidized beyond ~QFM + 1. If these processes had occurred, the solidification of the second magma ocean could have led to compositionally unique igneous rocks. Sulfide and chloride-rich melts could be separated from silicate magmas and crystallized alone.  During a lesser runaway greenhouse heating during the post-aqueous period (e.g., Kasting 1988), only some crustal material could have been metamorphosed via dehydration and decarbonatization. Some materials (e.g., chloride evaporites, felsic rocks) could have been melted, mobilized, and  \nredeposited. Sulfates and sedimentary sulfate-rich rocks could survive melting and assimilation.  \nDespite the global volcanic/tectonic resurfacing ( Sects. 16.2.2, 16.4.4.1 ), some morphological, mineralogical, chemical, and isotopic signs on post-aqueous metamorphic and igneous processes could be preserved on Venus' surface. Ejecta from several large impact craters (e.g., Mead, Stanton, Boleyn) with an elevated high dielectric constant (Pettengill et al. 1992) could inform on the composition of ancient rocks at a depth of several km. The composition of layered rock formations in tessera plateaus, canali, and outflow channels ( Sects. 16.2.2 ) may provide insights into past environments relevant to sulfates, sulfides, and other phases (chlorides, ferric oxides, silica minerals) and rock formations such as BIF's analogs formed in the presence of liquid water. Sulfates, phosphates, silica, and ferric oxides in sediments (or as secondary phases) supposedly deposited from aqueous solutions are stable on the current surface. Chlorides, carbonates (except MgCO3), and chlorides are subjected to sulfatization by SO2 ( Sect. 16.3.2 ) and may not be remotely detected.", '16.4.3 Volcanic Degassing, Space Sources, and Mass Balances': "Roughly compatible masses of C and N in the outer shells of Earth and Venus (Table 16.12) suggest similar degassing pathways. However, higher relative abundances of 20 Ne, 36 Ar, and 84 Kr, and the lower mass of radiogenic 40 Ar in the atmosphere of Venus imply differences in degassing among the planets (Avice et al. 2022; Gillmann et al. 2022, for reviews). The lower 40 Ar content (25-30% of Earth's) suggests smaller and/or earlier degassing of its interior. This alignment in scale but earlier degassing is consistent with the atmospheric abundances of C, N, and 40 Ar accumulated via decay of 40 K by the time of degassing.  This could reflect an early thorough degassing of the mantle. Another possibility for limited degassing could reflect the establishment of the present stagnant lithosphere early on in history (O'Rourke and Korenaga 2015; Rolf et al. 2022).  \n<!-- image -->  \nThe sulfur inventory within Earth's atmosphere-hydrosphere-crust system suggests volcanic degassing as the primary source of sulfur, rather than sulfide weathering, and sulfur is considered an 'excess volatile' in this system (Rubey 1951) alongside N, C, Cl, and H2O. Rubey (1951) noted the similarity between relative bulk compositions of 'excess' and volcanic volatiles. Although terrestrial volcanic gas sulfur is mainly trapped in seawater and sulfate minerals, biological sulfate reduction and subduction of oceanic lithospheric plates (Bekaert et al. 2021) confine assessing the mass of degassed sulfur throughout Earth's history. Nevertheless, subducted slab-delivered fluids are depleted in 34 S (δ 34 S = -2.5 ± 3%), suggesting a negligible amount of subducted sulfates (Li et al. 2020) and the current abundance of non-magmatic sulfur (mainly in sedimentary and oceanic sulfates) may roughly constrain the amount of degassed sulfur, as it was done by Rubey (1951).  \nThe comparable abundances of nitrogen and carbon in Earth's outer envelopes and Venus' atmosphere may suggest the existence of an 'excess' mass of sulfur in Venus' interior resulting from volcanic degassing. Table 16.12 depicts estimates of the amount of 'excess' Venus' sulfur from C, N, and S masses in terrestrial materials and Venus' atmospheric abundances. These assessments indicate that the possible molar amount of outgassed sulfur on Venus is an order of magnitude less than nitrogen and two orders less than carbon. However, Venus' atmosphere has much higher N/S and C/S ratios than assessed. Compared to the estimated amount of degassed sulfur, the negligible mass of Venus' atmospheric sulfur implies a substantial reservoir of trapped sulfur within the crust and/or in subducted materials. The prevalence of igneous rocks on Venus' surface ( Sect. 16.2.2 ) implies a significant capture of degassed sulfur before the global volcanic resurfacing event occurred 0.3-0.9 Ga ago. Trapped sulfur could be stored in igneous rocks containing assimilated crustal materials, sedimentary rocks formed via aqueous ( Sect. 16.4.2 ) or anhydrous (e.g., aeolian) deposition, and metamorphic (Harlov 2024) formations. Investigating the abundance, mineralogy, and isotopic composition of sulfur in these rocks could provide insights into past atmospheric conditions (such as via S-MIF) and into the environmental context (aqueous vs. anhydrous, aqueous oxidizing vs. aqueous reducing) during the trapping of atmospheric sulfur of volcanic origin, magmatic accumulation of crustal materials, and on possible subduction of S-bearing materials to the mantle.  \nThe amount of S-bearing species in a volcanic gas depends on the solubility and initial abundance of sulfur in magma. The temperature and pressure enhance the solubility, and the magma composition has complex effects, primarily reflecting the redox state and abundance of ferrous iron (Wallace and  \n<!-- image -->  \nCampbell 1992; Oppenheimer et al. 2011; Boulliung and Wood 2022, 2023; Simon and Wilke 2024; Casas  \net al. 2024). The mafic composition of Venus' volcanic rocks and oxidation state at landing sites ( Sect. 16.2.2.1 ) suggest sulfur solubility comparable to tholeiitic melts. Pressure is the main factor that affects the exsolution of S-bearing gases and determines the amount of degassed sulfur. Although Venus' atmospheric pressure (~47-110 bars at different elevations, Seiff et al. 1985) suppresses the degassing of silicate melts, low-solubility species (CO2, CO, H2, N2) are preferably degasses compared to highsolubility H2O and SO2, as first noted by Holloway (1992). Gaillard and Scaillet (2014) modeled the gasmelt separation of supposed Earth-like (CO2, H2O, S) volatiles in basaltic magma as a function of vent pressure and inferred low (SO2 + H2S)/CO2 and H2O/CO2 gas ratios at 10 2 bars that could characterize volcanic vents on current Venus. The S/C atomic ratio ~3 × 10 -3 was assessed for basalt-sourced gases at the QFM f O2 buffer and ~100 bars, while the initial magmatic ratio was ~5. However, experiments on SO2 solubility and models of Boulliung and Wood (2023) for basaltic melts suggest degassing at least 80% sulfur before the ascending magma reaches 10 3 bars pressure. Comparable amounts of CO2 and SO2 in basaltic gases at ~100 bars are also inferred with the comprehensive model of Ding et al. (2023) for sulfur degassing in CO2-H2O-S-bearing volcanic systems. For the Kilauea 1918-1918 lava lake gas composition (Gerlach 1980) with S/C of 0.2-0.8, Ding et al. (2023) suggest their release from pressures below 100 bars. For 1 bar, their model predicts the same S/C ratio in the gas as in parent magma. These results imply a thorough degassing of SO2 and the volcanic gas S/C ratio close to pre-eruptive values in basaltic melts for Venus' vents.  \nZolotov and Matsui (2002) calculated gas phase thermochemical equilibria to model the speciation of possible Venus' volcanic gases as a function of the C/H/S/Cl ratios and f O2. The nominal bulk gas composition was based on Kilauea 1918-1918 lava lake gas composition (Gerlach 1980) with a comparable amount of S and C (S/C of ~0.5). Their models exhibiting low H/C ratios echoed the current atmospheric composition ( x H2O/xCO2 = 3 × 10 -5 /0.965) and reflected the greater high-pressure solubility of H2O relative to CO2 (Fig. 16.28). At the QFM f O2 buffer, modeled H-poor gases are rich in CO2, and Sbearing gases are SO2, S2, S2O, OCS, S2Cl, SSF2, S3, SO, SCl2, and S2Cl2. At the IW buffer, H-deficient gases comprise CO, OCS, S2, CS2, and S3 and encompass a variety of S-halogen species. S-bearing species in volcanic gases, where H/C > 10 -2 , are primarily denoted by SO2, S2, H2S, S3, and SO at the QFM buffer and by OCS, S2, H2S, and S3 at the IW buffer. H-poor Venus' volcanic gases with comparable S and C atomic abundances mainly consist of CO2, SO2, CO, OCS, and S2O (Fig. 16.29). Gases with a low S/C ratio, as  \n<!-- image -->  \nproposed by Gaillard and Scaillet (2014), consist of CO2, CO, and SO2, with SO2 strongly dominating other sulfur-bearing gases.  \nThe near-surface atmosphere's supposed bulk and chemical composition (Table 16.1) strongly differ from volcanic gases. Assuming atmospheric origin via degassing of the interior, higher atmospheric C/(S, Cl, F) ratios imply trapping of degassed S, F, and Cl to minerals and the lack of carbonates in the crust (c.f., Zolotov 2018). The compositional difference between volcanic and near-surface gases (Fig. 16.30) also suggests post-eruption gas-phase thermochemical reactions that diminish the abundances of SO2, OCS, S2, S2O, CS2, Sn, and other volcanic gases. The difference would allow distinguishing volcanic plumes with remote or in situ methods (Wilson et al. 2024) to constrain current volcanic activity and physicalchemical conditions in magma and volcanic vents.  \nThe vast volume of mafic lava assessed from the mapping of Magellan radar images (140-250 Mkm 3 , Ivanov and Head (2013), (4.1-7.3) × 10 20 kg at 2900 kg m -3 density) implies the release of S-bearing gases during a geologically short period of widespread volcanic activity. Head et al. (2021) estimated that ~1/3 of current atmospheric sulfur could have been degassed from corresponding melts.  Their estimations were based on the model of Gaillard and Scaillet (2014) when 16 ppmw S degassed from magma with an initial 1000 ppmw S content. More S could be degassed using the model of Boulliung and Wood (2023) mentioned above. If 10-100% S is degassed from 800-1700 ppmw S in initial magma, 3 × 10 16 to 10 18 kg sulfur could have been released from 140-250 Mkm 3 basaltic melt.  This amount is comparable with the atmosphere's sulfur mass and the possible mass of 'excess' sulfur in the atmosphere and crust (Table 16.12) supposedly trapped after volcanic degassing supposedly trapped before the last global volcanic activity. In contrast to sulfur, similar abundances of S and C in volcanic gases do not suggest a significant contribution from the most recent global volcanic activity to the atmospheric CO2 inventory, aligning with the estimations of Head et al. (2021) and López et al. (1998). It is plausible that current atmospheric S-bearing species are products of global volcanic degassing that were partially trapped to S-bearing minerals in the surface materials ( Sect. 16.3.2.2 , Table 16.8). Even if volcanic activity occurs in the current epoch, these estimates suggest that the abundance of sulfur in the atmosphere may not be significantly impacted by such activity. The supply of reduced S-bearing gases released through the oxidation of exposed sulfides (e.g., Eqs. (16.80), (16.81)) could be more critical than current volcanic degassing.  \n<!-- image -->  \nSpace sources of sulfur occurring mainly in cometary dust could provide only a tiny fraction of the atmospheric sulfur within the last 0.5 Ga (Table 16.12). The resistance time of space-delivered sulfur obtained by the division of atmospheric sulfur mass by today's flux is 70 ± 41 Ga.", '16.4.4.1  The Global Volcanic Resurfacing': 'The Magellan radar data (Figs. 16.8-16.13) suggest any evidence of early geological history before 0.30.9 Ga (McKinnon et al. 1997; Korycansky and Zahnle 2005; Herrick et al. 2023) was obliterated due to global tectonic and volcanic events over a relatively short geological timescale. These processes resulted in the formation of tessera plateaus, volcanic plains, and volcanic-tectonic structures, such as ridges, volcanic centers, and coronas ( Sect. 16.2.2 , Strom et al. 1994; Basilevsky et al. 1997; Ivanov and Head 2011, 2013; Ghail et al. 2024). While the possibility of current volcanic activity is not ruled out (see Herrick et al. 2023 for a review), most volcanic rocks observed were emplaced during global resurfacing, and surface modification has been marginal since then (Carter et al. 2023). The degassing of S-bearing gases during that brief period likely contributed significantly to the current atmospheric sulfur ( Sect. 16.4.3) . Following the event, a considerable amount (potentially 50 ± 40%) of the degassed sulfur was captured in surface materials. The trapping is substantiated by the abundant sulfur found in surface samples (Table 16.3), the favorable reactions of SO2 with basaltic glasses and Ca-rich pyroxenes ( Sect. 16.3.2.2 ), and the equal masses of sulfur in the atmosphere and within a global 2-6 m thick layer of partially altered surface materials containing 1-3 wt% S (c.f., Lewis and Kreimendahl 1980; Zolotov and Volkov 1992). These estimates indicate a moderate reduction in the mass of atmospheric sulfur, the mixing ratios of S-bearing species, and changes in the properties of clouds since the global volcanic resurfacing.  \nThe models propose SO2 as the primary S-bearing volcanic gas on Venus, at least during the last global volcanic activity ( Sect. 16.4.3 ) at f O2 between the QFM and IW buffers ( Sect. 16.2.2.1 ). During the period of global volcanic resurfacing, the increase in atmospheric f SO2 facilitated reactions between SO2 and solids ( Sect. 16.3.2.2 ). Greenhouse heating due to degassing of H2O, SO2, and CO2 (e.g., Bullock and Greenspoon 2001; Solomon et al. 1999; Warren and Kite 2023; Way et al. 2016) increased rates of SO2solid reactions. Although SO2-solid reactions could have formed an array of minerals, the prompt  \n<!-- image -->  \nformation of Ca and Na sulfates ( Sect. 16.3.2.2 ) was unavoidable. The varying degrees of sulfur trapping observed in surface probes (Tables 16.3, 16.8) suggest efficient SO2-solid reactions during global volcanic activity. Volcanic materials placed earlier, possibly exemplified by Vega 2 rocks, could have trapped more sulfur than those at the Venera 13 (Fig. 16.1) and especially the Venera 14 site that exhibits less physically altered rocks (Zolotov and Volkov 1992). Morphological analysis of Magellan images and geological mapping (Abdrakhimov 2005; Basilevsky et al. 2007; Ivanov and Head 2011; Weitz and Basilevsky 1993) have shown that landing regions of Venera 13, 14, and Vega 2 are presented by the widespread type of volcanic plains (e.g., plains with wrinkle ridges) formed during global resurfacing. The location of these plains in the middle of the stratigraphic sequence (Ivanov and Head 2011) implies a higher degree of sulfur trapping in units placed before them. This suggests that these preceding units could be more altered than the S-rich Vega 2 rock.  \nGilmore et al. (2015) proposed that tessera plateaus formed shortly before a significant phase of plainforming volcanism (Ivanov and Head 2011; Kreslavsky et al. 2015) may have undergone ample alteration in atmospheric environments caused by widespread volcanic activity. They suggested that formed alteration products contribute to the low near-IR emissivity of tesserae surfaces ( Sect. 16.2.2.4 ). Sulfatization of highlands during the formation of volcanic plains and structures is consistent with an elevated near-IR emissivity of western Alpha Regio (Gilmore et al. 2015).  This region might have formed once the surrounding volcanic plain (Gilmore and Head 2000) was exposed to an atmosphere depleted by SO2. Consequently, it experienced less alteration by SO2 than the older regions of Alpha tessera. As is in the present epoch ( Sect. 16.3.2.2 ), lower temperatures at highlands enhanced thermodynamic affinity for sulfatization.  The occurrence of sulfate-rich materials with low thermal emissivity on highlands is a firm hypothesis that upcoming missions can test ( Sect. 16.5.2 ). The degree of sulfatization of geological formations estimated with remote sensing methods would constrain relative times of placement of surface materials.', '16.4.4.2  Mineral-Gas Buffers and Atmospheric Evolution': "Current conditions on Venus' modal radius are close to those determined by the Mt-Hem, Mt-Py, or MtPy-Hem mineral buffers ( Sect. 16.3.2.3, Figs. 16.22, 16.25, Table 16.11) that are used to control fugacities of O- and S-bearing gases in petrological and hydrothermal experiments. As Mueller proposed (1963, 1964, 1965), some assemblages could control trace atmospheric gases' fugacities. Despite some  \n<!-- image -->  \naeolian activity (Kreslavsky and Bondarenko 2017; Carter et al. 2023), the surface modification rate since the global resurfacing (~2 × 10 -3 µm per year, Arvidson et al. 1992) was orders of magnitude lower than on Earth.  This allowed atmospheric gases to interact with the same permeable materials over several hundred million years, favoring gas-solid type equilibration. The larger masses of Fe- and S-bearing minerals in permeable materials (at least tens of meters) compared with that of trace gases (SO2, OCS, CO, S2, etc.) support the feasibility of the buffering.  As in petrological experiments, the magnetite-tohematite transition would buffer the redox state of near-surface air that cannot become more oxidized than that at the Mt-Hem equilibrium until after the oxidation of all ferrous iron in exposed crustal materials. The latter is unlikely because of a limited H escape to space from the upper atmosphere (Chaffin et al. 2024; Gillmann et al. 2022 for a review), abundant ferrous iron in widespread and thick basaltic formations, and likely kinetic inhibition of the magnetite-to-hematite transition at nearequilibrium conditions. The decent match of atmospheric conditions with those controlled by major likely products of chemical weathering (pyrite, magnetite, hematite, anhydrite, Sect. 16.3.2.3 ) is unlikely to be accidental. It could result from the physical-chemical evolution of the coupled atmosphere-surface layer system following global resurfacing (Zolotov 1992a, b, 1995a, b, 2015).", 'Possible Formation of Pyrite': "Compared to sulfates ( Sect. 16.4.4.1 ), the alteration and formation of sulfides by gas-solid reactions during the global volcanic resurfacing are less noticeable. The elevated temperature favored the oxidation of ferrous and sulfide phases by CO2 (e.g., (16.80)). In addition to contributing to greenhouse heating, degassed water vapor oxidized Fe(II) in silicates (Eq. (16. 87)) and sulfides in volcanic products. As in the previous epochs, photochemical dissociation of H2O and hydrogen escape could have contributed to oxidation (Warren and Kite 2023). The high temperature favored the formation of Fe(III) oxides via oxidation at terresa terrains and other highlands.  \nPyrite is more stable in lower-temperature highlands, and ferric oxides are more stable in hot lowlands ( Sect. 16.3.2.3 , Klose et al. 1992; Zolotov 2018, 2019).  At constant temperatures, higher f SO2 favors the stability of pyrite, magnetite, or hematite, depending on the fugacities of other gases (S2, CO, O2) (Fig. 16.22, Zolotov 2018, 2019). Despite the temperature and fugacities remaining indifferently constrained for the volcanic resurfacing period (Solomon et al. 1999; Warren and Kite 2023).  The heating favored the decomposition of any pre-existing pyrite to S2 and pyrrhotite (Eq. (16.76)) that could be oxidized to iron oxides ( Sect. 16.3.2.3 ). The formation of secondary magnetite and hematite was more plausible  \n<!-- image -->  \nthan that of pyrite, particularly at significantly elevated temperatures that let crystal structures be relaxed (Solomon et al. 1999).  \nSecondary pyrite could have formed in highlands exposed to an SO2-rich atmosphere that had cooled down after the cessation of volcanic degassing and the decline of atmospheric H2O via rapid gas-solid reactions and hydrogen scape. In addition to the interaction of SO2 with ferrous compounds, pyrite could have formed via pyrrhotite oxidation by CO2 (Eq. (16.80)) and H2O. These conditions could have occurred during a late stage of volcanic activity and in a subsequent period. Pyrite formed in this manner could contribute to the low microwave emissivity observed in highlands ( Sect. 16.2.2.5 ).  \nGiven the current conditions of f SO2 and temperature and their proximity to the pyrite stability field ( Sect. 16.2.2.3 ), it is conceivable that pyrite could have formed on volcanic plains at elevated f SO2, particularly as the temperature significantly dropped towards the current values (~730-745 K). The potential presence of high-dielectric-constant material in the shallow subsurface (below 0.7-0.9 m, Antony et al. 2020) on these plains could suggest the formation of secondary pyrite on lava during or shortly after global volcanic activity. The low dielectric constant in the upper decimeters of plains (Pettengill et al. 1992) suggests fine-grained sedimentary deposits atop altered lava containing pyrite and/or iron oxides. This layered and porous material, observed at the landing sites of Venera 13 (Fig. 16.1) and 14, could be represented by lithified fine-grained deposits (Florensky et al. 1983) resulting from impacts (Basilevsky et al. 2004), explosive eruptions (Garvin et al. 1984; Basilevsky et al. 1985), wind action (Kreslavsky and Bondarenko 2017; Bondarenko and Kreslavsky 2018), and other sedimentary processes reviewed by Carter et al. (2023) ( Sect. 16.2.2.1 ). Note that a high-dielectricconstant material in shallow subsurface could be magnetite formed with Fe(II) oxidation by trapped water upon crystallization of basalts (Holloway 2004) and the low microwave emissivity could reflect volume scattering (Campbell et al. 1999; Tryka and Muhleman 1992) of rough subsurface materials rather than composition. The low emissivity of several large impact craters (Mead, Boleyn, etc.), volcanic edifices, and some tectonically formed measures (e.g., ridges) reported by Pettengill et al. (1992) agrees with the latter explanation.  \nWood (1994, 1997) suggested that pyrite could have formed through SO2-solid reactions on plains at 680-720 K (as opposed to the current 740 K) and in low-temperature highlands before global volcanism and associated SO2 degassing and greenhouse warming. According to Wood, weathering reactions  \n<!-- image -->  \nfollowing the global volcanic event and the partial consumption of SO2 did not result in pyrite formation on plains. Wood's hypothesis aligns with materials with high dielectric constants in the ejecta of large impact craters, the high microwave emissivity of volcanic plains, and the low emissivity noted on highlands (Pettengill et al. 1992). However, the survival of pyrite is unlikely in the case of severe greenhouse heating during global volcanism. If pyrite did not survive, the high-dielectric-constant in the deep surface could be presented by Fe(III) oxides formed via oxidation at gas-magma ( Sect. 16.4.1 ), gassolid, or aqueous ( Sect. 16.4.2 ) settings. Additional data ( Sect. 16.5 ) are required to understand the potential formation of iron sulfides and oxides before, during, and after the global volcanic resurfacing event.", '16.4.4.3 Possible Evolution Pathway to the Present and the Future': "Below is a probable pathway of the atmosphere-surface system since global resurfacing. A fraction of ferric oxides formed via high-temperature oxidation of volcanic materials during global volcanism (Warren and Kite 2023) remained in the surface material. The atmosphere was cooled via the partial consumption of degassed H2O and SO2 via oxidation and sulfatization (Eqs. (16.62), (16.63), (16.71), (16.74), (16.75), (16.87)). Hydrogen escape (Chaffin et al. 2024; Gillmann et al. 2022) that followed photochemical and chemical (Eq. (16.87)) reactions of water vapor provided net oxygen for further Fe(II) oxidation in silicates, glasses, and sulfides. Once hematite and magnetite formed in surface materials, net oxygen (in CO2, H2O, and SO2) was consumed by forming magnetite rather than hematite. The MtHem equilibrium sat oxygen fugacity in the near-surface gas phase and the pore spaces (Eqs. (16.65), (16.66)). Like the laboratory buffer, the ratio of hematite to magnetite changed as reactions took place; f O2 remained set by the buffer but varied with temperature. The redox state of the near-surface gas could not have become more oxidized than that at the Mt-Hem equilibrium until the ferrous iron in a permeable rock layer was oxidized. The latter was unlikely because of the suppressed volcanic degassing of H2O at 10 2 bars total pressure (Gaillard and Scaillet 2014), limited atmospheric water abundance, and abundant Fe(II)  in widespread and thick (Ivanov and Head 2013) basaltic formations. Given that O2 was virtually absent in the near-surface atmosphere, the Mt-Hem assemblage established the mixing ratios of SO2 to COS, H2S, and S2, as well as CO2 to CO and COS. The surface temperature reached a minimum and steadily increased with increasing solar luminosity. Correspondingly, the lower atmospheric f O2 at the Mt-Hem equilibrium initially decreased, only to increase subsequently.  \n<!-- image -->  \nAs in the preceding time, pyrrhotite was oxidized to magnetite by CO2 ( Sect. 16.3.2.3 ) and H2O. This oxidation could have transpired along the Mt-Pyrr phase boundary (Figs. 16.22, 16.25), potentially influencing the chemistry of the near-surface atmosphere. As the atmospheric temperature dropped, pyrite began to form on the highlands through reactions (16.78) to (16.81). The lowest temperature (before the increase) could be conducive to global pyrite formation and might account for its presumed presence on plains below ~0.7 m, as Antony et al. (2022) suggested. Following pyrite formation on the cooled plains, the atmosphere's and surface's linked evolution could have occurred along with the MtPy-Hem buffer (16.83), (16.84) (c.f., Zolotov 1992a, 1995a). Swift gas interactions with Mt-Py-Hem mixtures and individual minerals, as suggested by Venus-focused experiments (Reid 2021; Reid et al. 2024; Santos et al. 2024), ensure gas-solid interactions and equilibration within the geological time scale. The Mt-Py-Hem assemblage controlled f O2 and f S2, along with trace chemically active gases' fugacities and mixing ratios, except for HCl and HF. The subsequent formation of iron oxides and pyrite (Eqs. (16.80), (16.87)) was limited by the hydrogen escape rate that could be much faster (Chaffin et al. 2024) than previously assessed. No significant hematite formation is anticipated at f O2 set by Mt-Py-Hem and Mt-Hem buffers because of zero thermodynamic drive at chemical equilibrium. However, the ongoing accumulation of hematite from cosmic sources (Mogul et al. 2024; Zolotov 2021; Zolotov et al. 2023) could have contributed to the buffering assemblage(s) and is consistent with the red spectral slope observed at Venera landing sites (Pieties et al. 2006).  \nAs highlighted in Sect. 16.3.2.3 , the ambiguity of the atmospheric composition, the existence of impure minerals in buffering assemblages (e.g., Ti-bearing specimens, Zolotov 1994, 1995a, b), and a substantial f O2 error bar (~0.8 log units) at Mt-Hem equilibrium (16.66) (Fegley et al. 1997) complicate the distinction between possible buffering assemblages. At a specific temperature, the Mt-Py-Hem equilibrium determines both f O2 and f S2, the Mt-Py equilibrium (Eq. (16.77)) sets the f O2/ f S2 ratio within a range between the Py-Pyrr and Mt-Hem buffers, and the Mt-Hem equilibrium controls f O2 (Fig. 16.25, Table 16.11). Establishing conditions for the Mt-Py-Hem equilibrium, which solely depends on the greenhouse temperature, is a logical progression of the ~0.3-0.9 Ga evolution of the atmospheresurface system that resulted in the current near-surface physical-chemical environments. The suggested evolution, which includes a later addition of pyrite to the Mt-Hem assemblage, is more appropriate than developing the Mt-Py equilibrium assemblage (Eqs. (16.67), (16.77)) and the atmospheric evolution along it. The establishment of buffering mineral assemblages implies changes in f S2, f O2, and  \n<!-- image -->  \ncorresponding trace gases (SO2, OCS, CO, etc.) along the corresponding equilibria as temperature changes.  \nBoth SO2 degassing and elevated greenhouse temperatures during the global volcanic activity favored the sulfatization of exposed Ca-and Na-carbonates and the thermal decomposition of crustal carbonates if they existed. If iron sulfide-oxide equilibria started buffering trace gases after that activity, lower temperatures corresponded to lower f O2 and f SO2/ f S2 ratios (Zolotov 1992a, 1995a, Fig. 16.31), which were less favorable for sulfatization. Zolotov (1995c) explored the effect of temperature, f CO2, and f SO2 buffered by the Mt-Py-Hem assemblage on carbonate stability. The models do not exclude stable Cabearing carbonates and their sulfatization if temperature and f SO2 increase toward preset environments. Whether carbonates became stable is unclear, and the inquiry requires comprehensive modeling.", 'Future of the Atmosphere and Crust': "This research topic is still in its early stages and was briefly discussed by Gillmann et al. (2022). The increasing greenhouse heating will influence the coupled evolution of surface mineralogy and atmospheric composition due to rising solar luminosity, the declining x SO2 through sulfatization, and hydrogen escape. If another global volcanic event occurs, physical-chemical processes would be like those during the last global resurfacing ( Sect. 16.4.4.1 ). Trace chemically active gases would remain buffered by minerals at increasing temperatures. As in the preceding epoch, the Mt-Py-Hem assemblages will likely play a significant role in determining concentrations of trace C-O-S-H gases, though an alternative evolution among Mt-Py buffer was discussed as an alternative (Zolotov 1992b). One would expect changes in CO2/CO, SO2/COS, SO2/H2S, and H2O/H2 ratios that reflect increasing f O2 and f S2 with temperature. Possible thermochemical evolution along the Mt-Py-Hem buffer (e.g., Eqs. (16.83), (16.84)) suggests increasing f SO2 and CO2/CO ratio and subsequent thermal decomposition of pyrite (Fig. 16.31). Without a volcanic supply of H2O, the atmosphere will be desiccated via hydrogen escape (Chaffin et al. 2024) and oxidation of ferrous iron by (16.87) in exposed rocks. Atmospheric Sbearing gases will be increasingly affected by the thermal decomposition of crustal sulfides that release S2. As in the present epoch, atmospheric SO2 will not allow the formation of Ca, Na, and K carbonates. CO2 and N2 will remain the major atmospheric gases. Then, the Sun will become a red giant, and significant greenhouse heating will cause melting and/or thermal decomposition of salts (e.g., 'oceanic' chlorides), sulfides, sulfates, and silicates. These events and corresponding melt-atmosphere  \n<!-- image -->  \ninteractions will reduce the atmosphere by magma that is rich in ferrous and sulfide melt complexes. One would expect decreased CO2/CO and SO2/(H2S, S2, COS) atmospheric ratios towards f O2 between QFM and IW buffers that characterize Venus' basalts. The gas -melt equilibria at the surface of a third magma ocean in Venus' history will control the composition of its atmosphere, in which CO2, CO, N2, SO2, OCS, and S2 could be significant gases.", '16.5    Outstanding Questions and Future Exploration': "Understanding questions about Venus' sulfur involves examining the abundances of S -bearing compounds and their roles in current atmospheric and geological processes throughout history. The limited data on S-bearing species, especially in the lower atmosphere and geological materials, hinders our understanding of the fate of sulfur. This information scarcity restricts the use of S-bearing species as indicators for atmospheric processes, atmosphere-surface interactions, and coupled geological and climatic evolution. To address these gaps, new measurements of S-bearing gases and solids are crucial for enhancing our comprehension. Looking ahead, significant progress in Venus' exploration is anticipated before 2035, thanks to three selected space missions (DAVINCI, Garvin et al. 2022; VERITAS, Smrekar et al. 2022; and EnVision, de Oliviera et al. 2018, European Space Agency 2021), along with several proposed projects such as Venera-D (Zasova et al. 2017) and Venus Orbiter Mission (Shukrayaan1) aimed at exploring the planet's atmosphere, surface, and interior. Leveraging a new generation of ground-based and space telescopes holds promise for refining our understanding of atmospheric and cloud composition.  Details on selected and prospective missions, along with telescopic observations and modeling efforts, can be found in publications by Glaze et al. (2018), Limaye and Garvin (2023), and Wiedemann et al. (2023). Table 16.13 outlines the plans and approaches to investigate the abundance and speciation of sulfur in atmospheric and crustal materials. In addition to upcoming observations, numerical and experimental modeling of atmospheric processes, gas-solid interactions, and processes within the deep and shallow interior -such as magmatism, volcanism, and metamorphism -can contribute to further progress. Advancements can also be achieved through new laboratory data on the physical-chemical properties of atmospheric, surface, and interior materials.  \n<!-- image -->", '16.5.1 Abundances and Pathways of Atmospheric Sulfur': "In the upper mesosphere, one challenge lies in understanding the sources of sulfur that seemingly lead to increasing mixing ratios of SO2, SO, SO3, OCS, CS, and CS2 above 80 kilometers. There are two potential sources of sulfur: one from beneath (such as sulfuric acid aerosol) and the other from space, possibly as a component of interplanetary dust ( Sects. 16.3.1.2 to 16.3.1.4 ). Further measurements of Sbearing compounds can provide insights into their contribution to the mesospheric composition. Investigating whether there is a correlation between space fluxes, plumes, and mesospheric composition could provide clues about the origin of sulfur. Observers can look for compounds becoming abundant above 80 km and an abundance of elements (Fe, Si, Mg, C, Ni, etc.) indicative of space delivery. Telescopic observations of Venus' optical flashes, likely caused by meteors (Blaske et al. 2023), could serve as a proxy for assessing the intensity of space fluxes. Considering that space sources are evident on Earth and Venus (Carrillo-Sánchez et al. 2022), it is essential to account for space materials' ablation and photochemical alterations in mesospheric models. In the coming years, data on the upper mesosphere could be obtained via telescopic observations in the mm and sub-mm wavelengths and solar occultation data in the near-IR range (VIRAL spectrometer on Shukrayaan-1 orbiter, Patrakeev et al. 2022). Beyond observations above ~75 km, the underlying factors driving both short- and long-term variability of S-bearing species within the mesosphere could be elucidated through flyby and orbital investigations of cloud tops, as discussed below.  \nThe primary challenges in cloud studies involve characterizing and understanding S-bearing compounds, including their speciation and abundance. Constraining species distribution within the clouds and subcloud hazes, both vertically and latitudinally, and understanding sulfur fluxes from the lower and upper atmosphere is crucial. Key topics of interest include (1) The cause of long-term variability of SO2 at the cloud tops, (2) The composition and origins of the UV-blue absorber in the upper clouds, (3) The composition of aerosol particle size modes, (4) The presence of non-H2SO4 aerosol species, such as Sx and metal sulfates, (5) The composition of solutes in sulfuric acid, for example, metal sulfate/bisulfate complexes, (6) The potential for Sn species formation above the clouds and the feasibility of Sn supply from the lower atmosphere, (7) The origins, fate, and sinks of possible metal sulfates in the clouds and hazes, (8) The possibility of S-bearing organic compounds, such as an acid-insoluble organic matter of cosmic origin. Potential approaches to tackle these topics comprise long-term global measurements of cloud top composition (S-, C-, Cl-, N-bearing gases, and H2O) from telescopes and orbital platforms,  \n<!-- image -->  \nremote and in situ investigations of cloud dynamics (such as convective plumes and meridional circulation), and direct measurements of gas and aerosol composition within cloud layers and hazes.  \nFortunately, many planned clouds and hazes observations are relevant to these tasks (Table 16.13). In addition to telescopic observations, the mission plans include gathering compositional data on SO2 and SO at cloud tops and upper cloud aerosol in the UV spectral range using the DAVINCI (flybys), EnVision, Shukrayaan-1, and Venera-D orbiters, and with the Venera-D Lander Module. The EnVision VenSpec-H spectrometer will provide data to constrain abundances of SO2 and OCS through spectral windows within 1-2.5 µm below and above the clouds. As in preceding observations (Table 16.1), SO2 and OCS will be probed at 30-40 km in the night side spectral windows near 2.4 µm. Prospective data from the Venera-D VIKA orbital near-IR spectrometer will constrain S-bearing compounds in clouds, and the thermal IR Fourier spectrometer (5-40 µm) will provide data on SO2 at 55-75 km. Abundances of Sbearing gases and sulfuric acid aerosol at 45-55 km will be retrieved from radio occultations with EnVision and VERITAS (Akins et al. 2023). Venera-D MM-radiometer will measure H2SO4(g) and SO2 within and below clouds. Sampling and analysis of aerosol particles are planned with a chromato-mass spectrometer onboard the Venera-D Lander Module. Baines et al. (2021) proposed analyzing cloud aerosols at 52-62 km with a mass spectrometer via a balloon-borne gondola instrument package. Besides observations, the enhancement of our understanding of sulfur in the middle atmosphere will be facilitated by the development of cloud models that consider new measurements, chemical kinetics, the interactions of photochemical and thermochemical processes, the physical chemistry of interactions between gas, liquid, and solid phases, material delivery from above and below the clouds, sinks of particulates, vertical and lateral dynamics, and global circulation.  \nIn the lower atmosphere, the comprehension of S-bearing compounds relies on data regarding the abundance of SO2, SO3, H2SO4, OCS, CO, Sn, and H2S gases varying with altitude and latitude. Essential measurements yet to be made include the mixing ratios of SO2, OCS, CO, and Sn from the cloud deck to the surface, along with data on H2SO4 and SO3 gases at ~30 -48 km altitude. This information will aid in constraining the consumption mechanisms of SO3 and OCS at 30 -40 km, evaluating the origins of any existing SO2 altitudinal and latitudinal gradients, enhancing understanding of such gradients for OCS and CO, revealing the role of CO in reactions impacting sulfur chemistry at various altitudes and latitudes, and contributing to global circulation knowledge. In the near-surface atmosphere, it is crucial to measure the concentrations of chemically active gases (CO2, SO2, OCS, CO, H2S, Sn) to assess the degree  \n<!-- image -->  \nof gas-gas type equilibration, examine the hypothesis of such equilibration occurring in the lowlands (Table 16.2), constrain the atmospheric redox state from CO2/CO and SO2/(COS, H2S, S2) ratios, and understand the pathways of gas-solid reactions as well as the probability of atmospheric gas buffering by mineral assemblages ( Sect. 16.5.2 ). In the early 2030s, corresponding measurements will be conducted using the Venus' mass spectrometer (VMS) and a tunable laser spectrometer (VTLS) onboard the DAVINCI descent sphere (Garvin et al. 2022). The VMS will analyze SO2, OCS, H2SO4, H2S, Sn, and H2O from the cloud deck to the surface. The VTLS measures SO2, CO, H2O, and DHO/H2O ratios. The 2% accuracy of VTLS measurements of SO₂ and OCS and VMS sampling every 0.11 km (depending on the altitude) will provide profiles of these gases in the lower atmosphere.  Data from the Venus Oxygen Fugacity (VfOx) sensor aboard the DAVINCI sphere could complement information on gases and help constrain the degree of chemical equilibrium in the near-surface atmosphere. Data from the VenSpec-H near-IR spectrometer onboard the EnVision orbiter will help assess concentrations of SO2 and OCS in the night side lower atmosphere. The Venera-D Lander Module will be equipped with a gas chromato-mass spectrometer (VCS), a tunable laser absorption spectrometer (ISKRA-V), and a UV spectrometer (DAVUS) to measure the composition of SO2, SO, OCS, CO, and H2O during the descent from 70 km to the surface. Besides measurements, efforts should be made to advance coupled thermochemical-photochemical models, including global and local eddy diffusion and atmospheric circulation. The next logical steps are developing 2D (e.g., meridional) and 3D chemical-transport models before obtaining data from upcoming missions. To make the modeling fruitful, experimentally determining rates of gas phase reactions that are important for key science questions, such as OCS consumption at 30s km or Sn production, is also necessary.  \nSulfur isotopes ( 32 S, 33 S, 34 S, and 36 S) in terrestrial atmospheric aerosols and gases (SO2, OCS) offer valuable information about the origins and processes that influence sulfur (Lin et al. 2018; Angert et al. 2019; Eldridge 2024). Evidence of mass-independent fractionation of sulfur in tropospheric sulfate aerosols provides insights into stratospheric photochemistry. The 32 S/ 34 S ratio in atmospheric SO2 and OCS implies mass-dependent fractionation on Earth and sheds light on past and current biological pr ocesses, fossil fuel combustion, and volcanic degassing. Similarly, sulfur isotopes in Venus' atmospheric samples will provide insights into photochemical processes, circulation and mixing, and the contribution of isotopically distinct sulfur sources from the planet's interior and space. The meager 32 S/ 34 S ratio deduced from the Pioneer Venus Large Probe data ( Sect. 16.2.1.1 ) might be inaccurate, and  \n<!-- image -->  \nspectroscopic measurements of 32 S, 33 S, and 34 S in SO2 and OCS by the VTLS instrument are scheduled for the DAVINCI mission. Independent data could be obtained using the ISKRA-V and VCS spectrometers on the Venera-D Lander Module, as well as from a mass spectrometer designed for the Aerosol-Sampling Instrument Package (ASIP), which is intended for a balloon mission (Baines et al. 2021).  \nAnother question pertains to the impact of current volcanic activity on the composition of atmospheric gases and aerosols. Wilson et al. (2024) have compellingly shown that additional remote and in situ data on atmospheric composition can indicate ongoing volcanic activity. Fig. 16.30 illustrates that sampling a volcanic plume could provide valuable insights, allowing us to constrain volcanic gas composition and its intricate interaction with atmospheric gases.", '16.5.2 Abundances and Pathways of Sulfur in Surface and Interior Materials': "Like Earth (e.g., Casas et al. 2024, Eldridge 2024, Simon and Wilke 2024), Mars (Franz et al. 2014; Franz 2024), the Moon and Mercury (Renggli et al. 2024), and moons in the outer solar system (Fegley and Lodders 2024), the abundance, speciation, and isoto pic composition of sulfur in Venus' surface and interior materials would reflect the interplay between endogenic and exogenic processes throughout history. Given the limited data on sulfur in surface materials (Table 16.3), any additional insights into the composition of surface and interior materials could significantly enhance our understanding of sulfur's abundance, speciation, and role in current and past geological processes.", '16.5.2.1   Atmosphere-Surface Interactions': "Essential inquiries concerning gas-solid reactions at the surface include the extent of sulfur intake and the occurrence and composition of secondary sulfates and sulfides in alteration products of rocks, minerals, glasses, and fine-grained materials throughout the surface and in buried altered rocks. This includes specific geological features and locations at various latitudes and altitudes. One aspect under consideration is the potential presence of abundant sulfides and/or sulfates in the highlands and whether these minerals formed during the current epoch or alongside the magma degassing during the past global volcanic resurfacing ( Sects. 16.4.3, 16.4.4 ). Information on the abundance and mineralogy of sulfur at highlands such as tesserae will provide insights into the reasons for the low microwave and near-IR emissivity of surface materials ( Sects. 16.2.2.4 , 16.2.2.5) . The secondary mineralogy of sulfur will  \n<!-- image -->  \nbe tentatively constrained through remote sensing from several spacecraft (Table 16.13). Iron and trace metal sulfides in surface materials could be inferred from high-resolution radar studies by the EnVision and Shukrayaan-1 orbiters that will constrain dielectric properties. Near-IR surface emissivity data will be collected globally, regionally, and locally from VERITAS, EnVision, Shukrayaan-1, Venera-D orbiters, and the DAVINCI flyby bus and the descent sphere. These overlapping observations will be utilized to identify the types of surface rocks and whether these rocks are coated with low-albedo materials such as sulfates or opaques like iron sulfides and oxides (Dyar et al. et al. 2020, 2021; Treiman et al. 2021; Helbert et al. 2021; McCanta et al. 2024). While major types of uncoated rocks (mafic vs. felsic) can be identified, the specific mineralogy may not be determined remotely. However, observations of the thermal emission in the near-IR range over several years will assist in detecting surface changes and identifying recent and ongoing volcanic activity.  \nIn situ studies of surface materials by landers and rovers can serve as ground truth for remote-sensing investigations. Measurements on surface missions include assessments of sulfur abundance from XRF, Alpha Particle X-ray Spectrometry (APXS), Laser-Induced Breakdown Spectroscopy (LIBS), and gamma-ray spectrometry with neutron activation, along with assessments of phase composition investigations with XRD (X-ray Diffraction), as well as Mössbauer, Raman, UV, and near-IR spectroscopy. Microscopic optical and Raman imaging would provide insights into the alteration texture of rock fragments, both altered and unaltered mineral grains. Investigations of dust-coated, brushed, and drilled samples would help determine the degree of alteration and whether there has been sulfur uptake or loss. Analyzing the 32 S/ 33 S/ 34 S isotopic ratios in secondary phases will constrain the involvement of atmospheric sulfur. Some of these measurements have been proposed for several NASA Venus Flagship Mission concepts. They could be performed with the proposed Venera-D Lander Module equipped with the XRF/XRD spectrometer, an APXS spectrometer, a Mössbauer spectrometer, and a laser ablation mass spectrometer (Table 16.13, Widemann et al. 2023).  \nAnother set of inquiries pertains to the direction of gas-solid reactions driven by chemical disequilibria, which depends on the composition of exposed materials, atmospheric composition, temperature, and pressure. DAVINCI data on the composition of the near-surface atmosphere will enhance evaluations of mineral stability and potential directions of weathering reactions (Garvin et al. 2022). Even in the absence of new data on near-surface environments, we can still make progress in evaluating the stability and reactivity of solid phases on the current surface (refer to Zolotov 2019, 2018 for details). Promising  \n<!-- image -->  \nresearch areas include investigating the stability of yet-to-be-explored sulfides, sulfates, and minerals that could undergo sulfurization. Consideration could be given to studying silicate solid solutions of rockforming minerals (pyroxenes, felspars, etc.) and environmental conditions suggested by climate and integrated atmosphere-interior models. Further experimental studies of gas-solid reactions could be focused on thermodynamically plausible interactions. Investigating the advanced stages of alterations and the effects of proximity to equilibrium conditions on reaction rates is essential. Many initial reconnaissance and exploratory experiments have been performed (Table 16.9), laying the foundation for focused investigations of solid phases and glasses. These focused investigations can advance the field by deriving data on reaction rates, rate-limiting steps, and mechanisms throughout different stages of alteration, which can then be incorporated into numerical models to assess the alteration. The preliminary experiments have also highlighted two key improvements for the design of weathering experiments: the use of chemically inert reaction vessels to maintain gas compositions (for instance, by regulating gas flows or using mineral buffers) and the detailed analysis of gaseous by-products from gassolid interactions. By enhancing control over, or knowledge of, the gas composition throughout an experiment, we can more effectively incorporate the findings into our comprehension of gas-solid reactions. In addition to experiments, insights could be obtained from terrestrial analog studies in volcanic (McCanta et al. 2014; D'Incecco et al. 2024) and metamorphic settings.", '16.5.2.2   Sulfur in Geological and Climatic Evolution': "Outstanding, bigpicture questions about sulfur in Venus' history ( Sect.16.4 ) are related to the formation and differentiation of the planet and its putative early evolution that might have involved a magma ocean, oxidation of atmospheric, crustal and mantle materials, plate tectonics, and a hydrosphere. More Venus-specific inquiries include the history of volcanic resurfacing and degassing and their effects on physical-chemical processes in the atmosphere and sulfur trapping to crustal materials. These topics could be addressed through investigations of bedrocks with remote sensing and in situ sampling and geophysical investigations of shallow and deep interiors via radar sounding, gravity, topography (Table 16.13), and possible seismic measurements.  \nUnderstanding the fraction of accreted sulfur sequestered in the core versus that remaining in the mantle remains an outstanding question in Venus' exploration. In situ measurements of sulfur content in unaltered mafic rocks could offer constraints on the mantle-core partitioning of sulfur, and elevated  \n<!-- image -->  \nsulfur content in such rocks may signify an S-enriched silicate mantle. Concentrations of chalcophile elements in basalts also indicate sulfur partitioning to the core. Another promising approach involves analyzing copper isotopes within a basalt to provide insights into the segregation of sulfide liquids during core formation (Savage et al. 2015). The gravity data from the VERITAS (Smrekar et al. 2022; Cascioli et al. 2023) and EnVision (Rosenblatt et al. 2021) missions will address critical questions about the interior structure, including whether the core is liquid or solid, the presence of a basal silicate magma ocean (O'Rourke et al. 2020), and may offer constraints on the distribution of sulfur within the interior. However, direct inference of core sulfur content may prove elusive, akin to challenges faced with other terrestrial planets (Komabayashi and Thompson 2024).  \nFirst-order questions regarding sulfur in magmatic systems encompass major rock types ( Sect. 16.2.2 ), such as tholeiitic basalts, alkaline mafic rocks, evolved silicate rocks, and potential non-silicate rocks in canali and crater outflows. In addition to exploring the abundance and mineralogy of sulfur in these formations, these inquiries encompass the occurrence of cumulative magmatic and/or associated (e.g., hydrothermal) sulfide deposits. Near-IR and microwave emissivity measurements will constrain igneous rocks' composition, and sulfide-rich secondary or primary formations could be implied. High spatial resolution radar images, while not directly assessing composition, could be used to evaluate lava viscosity or reveal patterns of physical weathering that depend on composition.  Radar sounding conducted by the EnVision and Shukrayaan-1 orbiters could infer the presence of massive sulfides at depth . New radar data are needed to verify the presence of abundant iron sulfides and/or oxides below ~0.9 m, as suggested by Antony et al. (2022)  \nQuestions surround whether sulfide sulfur underwent oxidation due to residual oxygen remaining after hydrogen escape. We seek to understand the scale and context of such oxidation. The specific mechanisms remain elusive if sulfur, iron, and other elements experience oxidation. Did this process occur at the top of a magma ocean, through later gas-solid interactions, or via aqueous processes near the planetary surface ( Sect. 16.4 )? These pathways could intertwine, creating a fascinating puzzle. Further inquiries arise regarding the fate of oxidized rocks. Did they accumulate in the upper crust or descend through the subduction of lithospheric plates? Perhaps they were submerged (e.g., ElkinsTanton et al. 2007) or buried by basaltic lavas throughout history. We also ponder whether aqueous processes led to the formation of sulfate-bearing rocks and whether these rocks are exposed. Our  \n<!-- image -->  \ninquiries extend to understanding when, where, and how sulfates underwent reduction and whether putative biotic sulfate reduction led to the fractionation of 32 S and 34 S.  \nRedox- and paleoclimate-related inquiries can be addressed through investigations of bedrocks within diverse geological formations. Attention could be devoted to ejecta from Mead, Stanton, and Boleyn impact craters that contain rock fragments from a depth of several km. The abundance, mineralogy, and oxidation state of sulfur (S-II), S(-I), S(VI), and other elements (Fe, V, Eu, Ce, U, Mn, Mo, etc.) in igneous rocks are crucial for constraining the interconnected magmatic and climatic history. As observed in surface rocks on Mercury (Nittler et al. 2018; Renggli et al. 2024), a few wt% of sulfide's sulfur in unaltered igneous rocks would suggest significantly reduced mantle rocks ( f O2 < 3 log units below the IW buffer, e.g., Namur et al. 2016) and a scarcity of water in Venus-forming materials. On the contrary, detecting sulfates in silicate and/or nonsilicate rocks would indicate oxidation of Venus' interior due to the accretion of abundant water and following hydrogen escape. The detection of igneous rocks with sulfate sulfur and abundant ferric iron -rather than sedimentary or metamorphosed chemical sediments like banded iron formations -hints at an anhydrous oxidation event involving crustal materials during a magma ocean stage.  \nThe occurrence of putative sedimentary, metamorphic, or hydrothermally altered rocks within tessera terrains ( Sect. 16.2.2.4 ) will yield additional constraints for our understanding of these processes. Layered deposits rich in sulfates would suggest the presence of O2 in the atmosphere-hydrosphere system that preceded or accompanied aqueous (possibly evaporitic) deposition. The composition of metal (e.g., Ca, Mg, Na, K) sulfates will offer constraints on conditions of aqueous deposition, such as salinity, temperature, and pH . Deposits rich in chlorides and/or sulfides (e.g., pyrite) but sulfate-free would suggest aqueous deposition under anoxic, possibly high-temperature environments. The presence of sedimentary sulfides in Venus' rock formations, the mineralogy of sulfides and trace metal content, along with the isotopic composition of sulfur and trace metals (Mo, Re, etc., Lyons et al. 2014), will help determine the characteristics of aqueous environments, redox conditions, and the fate of sulfur on land and in water ocean. Sedimentary sulfide would also assess the likelihood of biological sulfate reduction. The MIF-S signature would suggest an anoxic atmosphere. Investigations of elemental, isotopic, and phase composition (Table 16.13) of at least two sites (plains and tessera regions) complemented by remote sensing data will be essential to elucidate the fate of sulfur during the interior-atmospheric evolution.  \n<!-- image -->", '16.5.3 Conclusion': "Our understanding of sulfur on Venus is constrained by the minimal data available below 20 kilometers of altitude in the atmosphere. Only one gas species (S3) has been tentatively detected in the nearsurface atmosphere, and the total sulfur content in surface materials has only been quantified with large error bars at three locations. However, the abundance of measurements of S-bearing gases at altitudes of 30 -120 km, the presence of thick sulfuric acid clouds, and the sulfur-rich composition of otherwise basaltic rocks all point to a significant role of sulfur in the atmosphere and reactions between the atmosphere and the surface materials. Due to sulfur's role in photochemical, thermochemical, volcanic, magmatic, mass transfer, and isotope fractionation processes, the abundance and speciation of sulfur in potentially detectable materials (gases, aerosols, surface solids) signify planet-wide processes currently and throughout history. For instance, elevated sulfur abundances in the atmosphere and surface probes imply past volcanic degassing of S-bearing species, which affected cloud formation and the trapping of atmospheric gases in minerals. If detected, sulfate and/or sulfide coatings on highlands would suggest sulfurization in the current epoch and/or during global volcanic degassing several hundred million years ago. The mineralogy and abundance of sulfur and chalcophile elements in pristine igneous rocks will provide insights into the fate of sulfur during planetary differentiation and igneous processes. Sulfates in igneous rocks would indirectly indicate the presence of abundant early water that affected oxidation of the mantle via hydrogen escape and/or formation and assimilation of sulfate-rich crustal rocks, such as evaporites. Similarly, detecting layered sulfate-bearing deposits in tessera terrains would suggest deposition in oxygen-rich aqueous environments, potentially indicating hydrogen escape from a moist atmosphere. The discovery of pyrite associated with iron oxides in surface materials would imply the buffering of sulfur-bearing gases and the redox state of the near-surface atmosphere in the current epoch. Such buffering suggests the preceding and future co-evolution of the coupled atmospheresurface system in which gas chemistry is influenced by volcanic degassing, resurfacing by lava flows, hydrogen escape, and increasing solar luminosity. Therefore, the study of S-bearing compounds should be a priority in the exploration of Venus. Continue d progress in our understanding of Venus' sulfur hinges on observations from terrestrial and space telescopes and flyby and orbital platforms. In situ data from atmospheric probes and landers is vital in deciphering the deep atmospheric chemistry, gas-solid interactions, and the destiny of sulfur in the interior. Significant limitations can be established by measuring the fundamental properties of S-bearing compounds and physicochemical systems and via  \n<!-- image -->  \nlaboratory modeling of physicochemical processes in the atmosphere, on the surface, and below it. Lastly, developing and applying updated and innovative numerical models are indispensable to link these efforts with observations of sulfur species and to align with other astronomical, geophysical, petrological, and geological information about the Earth's sister planet.  \nAcknowledgments. This chapter benefits from comments provided by Emmanuel Marcq, Kevin Zahnle, Larry Esposito, Alison Santos, Alexander Basilevsky, Michael Way, Laura Schafer, Lindy Elkins-Tanton, Ekaterina Kisseva, Molly McCanta, Joseph O'Rourke, Nancy Chabot, Rakesh Mo gul, Sara Port, and Julia Semprich. Over the years, my work on Venus was inspired and supported by Igor Khodakovsky, Vladislav Volkov, Alexander Basilevsky, and Bruce Fegley, Jr. The NASA Discovery and Solar System Workings programs supported the work on the chapter.", 'REFERENCES': "Abe Y, Matsui T (1988) Evolution of an impact-generated H2O -CO2 atmosphere and formation of a hot proto-ocean on Earth. J Atmos Sci 45:3081 -3101. https://doi.org/10.1175/15200469(1988)045<3081:EOAIGH>2.0.CO;2  \n- Abe Y, Abe-Ouchi A, Sleep NH, Zahnle KJ (2011) Habitable zone limits for dry planets. 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Lunar Planet Sci Conf 54:2880. https://www.hou.usra.edu/meetings/lpsc2023/pdf/2880.pdf  \n191  \n<!-- image -->  \nTable 16.1 Mixing ratios of S-bearing gases and CO in the atmosphere of Venus based on remote sensing and in situ data  \n192  \n<!-- image -->  \n<!-- image -->  \n194  \n<!-- image -->  \nPVPL: Pioneer Venus Large Probe  \nLNMS: Large Probe Neutral Mass Spectrometer  \nGC: gas chromatography  \n'ground' : telescopic observations from Earth  \nppmv: part by million by volume  \nppbv: part by billion by volume  \npptv: part by trillion by volume  \n1 CO data does not list all measurements  \n<!-- image -->  \nTable 16.2 Mixing ratios of chemically active gases in the near-surface atmosphere of Venus according to gas-phase chemical equilibrium models 1  \n1 Chemical equilibrium among near-surface atmospheric gases remains to be hypothetical. Concentrations in parentheses are based on instrumental data (Table 16.1) and used as anchors to assess the mixing ratios of other gases.  \n- 1: Krasnopolsky and Parshev (1979)\n- 2: Krasnopolsky and Parshev (1981a)\n- 3: Barsukov et al. (1980a), H2S value looks incorrect and omitted\n- 4: Barsukov et al. (1982), Dorofeeva et al. (1981), Khodakovsky (1982)\n- 5: Fegley et al. (1997b), the Gibbs free energy minimization method based on JANAF data; the result is like that of Zolotov (1996)\n- 6: Rimmer et al. (2021) Gibbs free energy minimization method at initial 96% CO2, 20 ppmv CO, 150 ppmv SO2, 5 ppmv OCS, 10 ppbv H2S, and 3 ppbv H2  \n7-10: this work. In 7, x SO2 is from Gel'man et al. (1980) and Marqc et al. (2021). In 9, the CO concertation of 8 ppm is obtained by extrapolating the CO gradient at 12 -36 km (Table 16.1) to the surface. In 10, x S3 corresponds to the data for 3 -10 km (Krasnopolsky 2013). In models 7 and 8, the gas composition resembles that the magnetite-pyrite buffer determines (16.66, Table 16.11). In models 9 and 10, the gas composition resembles that determined by 16.66 at f O2 controlled by the magnetite-hematite buffer (16.65, Table 16.11). Before obtaining new data for the near-surface atmosphere, models 8 and 9 could be considered as nominal reduced and oxidized cases, respectively  \n<!-- image -->  \nTable 16.3 Chemical composition of Venus' surface materials at the Venera and Vega landing sites (wt%) 1  \n1 Surkov et al. (1984, 1986b). The data were obtained with X-ray fluorescent analysis that does not provide speciation. Uncertainties are  1σ. Sulfur concentration is presented in the oxide and elemental forms. All Fe is given as FeO. Na has not been measured.  \n197  \n<!-- image -->  \nTable 16.4 Potassium, uranium, and thorium contents surface materials in sites of Venera and Vega landers 1  \n<!-- image -->  \nTable 16.5 Sulfur contents in models of Venus interior composition (BVTP 1981)  \nVe1: Equilibrium condensation models of the solar nebula. The absence of S in Ve1 is accounted for by an elevated nebula temperature at the radial distance where Venusforming materials condensed.  \nVe2: Equilibrium condensation models of the solar nebula that included feeding zones  \nVe3: Morgan and Ander's (1980) model is based on chemically modified nebular condensates that resemble stony and iron meteorites. K/U ratio from Mars 5 data as a K/U proxy for volatile/refractory elemental ratios  \nVe4: Pyrolite models of Ringwood developed based on analogs of Earth's model  \nVe5: an iron-deficient model  \n<!-- image -->  \nTable 16.6 Major sources and sinks of S-bearing species in the atmosphere on Venus 1  \n1 Additonal cells in 'sources' and 'sinks' designate atmospheric regions. S, surface and near-surface (< 10 km); LA, lower atmosphere; MA, middle atmospheres (clouds); UA, upper atmosphere (mesosphere)  \nTable 16.7 Major chemical reactions of S-bearing gases in the atmosphere of Venus 1  \n201  \n<!-- image -->  \n1 Numbers of the most important reactions are in bold. Values in parenthesis correspond to mean altitudes of reaction yield from Krasnopolsky (2012, 2013)  \n<!-- image -->  \nTable 16.8 Estimated degree of sulfatization of solid materials sampled at the landing sites of Venera 13, Venera 14, and Vega 2 1  \n1 The estimations are based on measurements from Table 16.3 and calculations of chemical equilibria in multicomponent gas-solid systems informed by the composition of solids at landing sites and the atmospheric composition on Venus (Barsukov et al. 1986c). Case (1) corresponds to the mineralogy of surface materials in equilibrium with an atmosphere open (fixed f gases) with respect to CO2, CO, H2O, SO2, HCl, and HF. The calculated sulfur (S) content reflects the trapping of atmospheric sulfur. Case (2) is for mineralogy equilibrated with the atmosphere open with respect to CO2, CO, and H2O. In this case, the S content corresponds to the measured values  \n2 The estimated amount of sulfate S in solid samples reflects 0.08 wt% S in magmatic sulfides, assumed by analogy with glassy parts of submarine erupted basalts (Moore and Fabbi 1971)  \n3 Degrees of sulfatization correspond to atomic ratios of supposedly sulfate S in analyzed samples to Ca and S in mineral assemblages equilibrated with atmospheric SO2. The higher degree of sulfatization in the last row reflects the presence of plagioclase (An20-30) in equilibrium assemblages and corresponds to the alteration of phases that are unstable concerning sulfatization (e.g., Ca-bearing pyroxenes)  \nTable 16.9 Experimental works relevant to interactions of S-bearing gases with geological materials at Venus surface conditions 1  \n204  \n<!-- image -->  \n1 Reactions in O2-bearing conditions are not listed  \nTable 16.10 Alteration products of minerals and glasses based on chemical weathering experiments at simulated temperature, pressure, and gas composition of Venus' surface 1  \n<!-- image -->  \n1 Experiments performed at 1 bar and with participation of one gas are not listed  \nTable 16.11 Mixing ratios and fugacity of chemically active gases at the conditions of the magnetite-pyrite equilibrium at 740 K and 95.6 bars 1  \n1 The gas content determined by the Mt-Py equilibrium (16.67, 16.77) is bracketed by fugacities and concentrations set by the Mt-Py-Pyrr (16.80, 16.82) and Mt-Py-Hem (16.83, 16.84) assemblages (Figs. 16.22, 16.23 -16.25, 16.31). Concentrations in parentheses are based on instrumental data (Table 16.1) and serve as anchors for assessing the concentrations and fugacities of other gases. Note the large uncertainties of gas abundances at buffers shown in the listed figures. The supposed Venus' surface conditions are within the un certainty of the Mt-Py-Hem buffer.  \n2 The gas composition corresponding to the Mt-Py equilibrium at 150 ppmv SO2 aligns with the in situ measurement of xCO at 16 ppmv (Table 16.1). This composition could be considered a nominal gas model for Venus' modal radius, consistent with other models (Fegley et al. 1997b, Table 16.2).  \n3 f S2 is set by the Py-Pyrr equilibrium 16.81 after Hong and Fegley (1998)  \n4 f O2 is set by the Mt-Hem equilibrium 16.66. Note the error bar of 16.66 of 0.8 log units (Fegley et al. 1997b)  \n<!-- image -->  \nTable 16.12 Abundances and sources of N, C, and S in the atmosphere-crust system of Venus  \n1,2,3 Data for N and C data for Earth and Venus are from Volkov (1992). For the Earth, sulfur values r epresent the 'excess' mass after Rubey (1951). The 'g/g' values are normalized to planetary masses 2 Venus' crust is assumed to be N - and Cfree. The values for sulfur correspond to the 'excess' amount calculated by equations SV-Ex = (NV /NE) × SE-ex SV-Ex = (CV/CE) × SE-ex where SV-Ex stands for the mass of 'excess' sulfur on Venus; N E and CE are masses of nitrogen and carbon in the terrestrial atmosphere and crust, NV and CV designate masses of N and C in Venus' atmosphere, and S E-ex is the mass of 'excess' sulfur in upper envelopes on Earth. These values represent sulfates and sulfides formed through the trapping of degassed sulfur throughout history  \n3 The mass of sulfur in the atmosphere of Venus corresponds to 150 ppmv SO2 (Volkov 1992)  \n4 The fluxes correspond to the flux of space material of (3.1 ± 1.8) × 10 4 kg/day (Carrillo-Sánchez et al. 2020). The sulfur values are assessed assuming the flux of cometary dust (~94% of the infall) containing 2/3 of C-, N-free inorganic mass fraction with 7.9 wt% S and 1/3 of organic (C100H80O30N3.5S3) mass fraction with 5.1 wt% S  \nTable 16.13 The planned and proposed investigations that could enhance our knowledge on sulfur and S-bearing compounds in the atmospheric, surface, and interior materials on Venus 1  \n<!-- image -->  \n1 Details can be found in Widemann et al. (2023), Smrekar et al. (2022), Garvin et al. (2022), Zasova et al. (2017), and European Space Agency (2021)  \n<!-- image -->  \nFig. 16.1 Layered rocks and rock fragments at the landing site of Venera 13. The original color images (Selivanov et al. 1983; Bokshteyn et al. 1983) were digitally mastered by Don Mitchel. The surface is black at Venus ' conditions, and the color represents the surface at room temperature (cf., Pieters et al. 1986). The composition of the material corresponds to potassium-rich mafic silicate rock with elevated sulfur content (Table 16.3)  \n<!-- image -->  \n<!-- image -->  \nFig. 16.2 SO2 mixing ratio at 70 km in the atmosphere of Venus as a function of latitude (Marcq et al. 2020). The red line represents a median value; the red bars show one statistical dispersion. The color pattern depicts % of observations. The data were obtained from the Venus Express SPICAV-UV observations (Table 16.1)  \n<!-- image -->  \nFig. 16.3 The long-term changes in SO2 mixing ratio at the cloud top of Venus' atmosphere (after Marcq et al. 2013 and Jessup et al. 2015). The blue cross represents the average value derived from STIS-HST observations  \n<!-- image -->  \n<!-- image -->  \nFig. 16.4 Latitudinal dependence of SO2 and CO mixing ratios at 33 km altitude (Marcq et al. 2021). The data were obtained from ground-based near-IR nightside observations  \n<!-- image -->  \n<!-- image -->  \nFig. 16.5 Abundance profile of H2SO4 vapor in lower and middle cloud regions (at 48 -55 km) and below clouds derived from the Magellan radio occultation data (Kolodner and Steffes 1998). Within clouds, the H2SO4(g) mixing ratio is consistent with the H2SO4(g) saturation curve (black curve). The decrease in x H2SO4(g) below clouds suggests its thermal decomposition to SO3 and water vapor  \n<!-- image -->  \n<!-- image -->  \nFig 16.6 Vertical structure of Venus clouds obtained with Pioneer Venus probes (Knollenberg and Hunten 1980)  \n<!-- image -->  \n<!-- image -->  \nVenera 9  \nVenera 10  \n<!-- image -->  \nFig 16.7 TV images of the surface in Venera 9 and 10 landing sites  \n<!-- image -->  \n<!-- image -->  \nFig 16.8 Magellan radar image of Venus volcanic plains (image by NASA/JPL). Lava flows are bright because of their rough surface. The imaged area is 140 km wide and located in the east part of the Sapas Mons  \n<!-- image -->  \n<!-- image -->  \nFig 16.9 Computer-simulated Magellan radar image of Venus volcanic center, Sif Mons (NASA/JPL). Sif Mons (upper part of the image) is a volcano ~300 km in diameter and a height of ~2 km. Magellan radar data is combined with radar altimetry, which is enhanced here. The hues are arbitrary and reflect colors recorded by Venera 13 (Fig. 16.1) and 14 landers. Like Earth's basaltic rocks, the actual surface is dark in the visible range  \n<!-- image -->  \nFig 16.10 Magellan radar image of steep-sided domes on Venus volcanic plains (NASA/JPL photo). The domes are ~25 km in diameter, ~2 km tall, and located east of Alpha Regio (~29.5° S, 12° E). They could have formed through the eruption of highly viscous silicate magma  \n<!-- image -->  \n<!-- image -->  \nFig 16.11 Magellan radar image of Baltic Vallis lava channel on Venus (NASA/JPL photo). The channel is ~2 km wide and shows branches and islands. The image is about 50 km wide. The location of the image is Sedna Planitia, south of Ishtar Terra (42.7° N, 340.7° E)  \n<!-- image -->  \n<!-- image -->  \nFig 16.12 Magellan radar image of a part of the northern boundary of Ovda Regio on Venus (NASA/JPL photo). The image is 300 km across. The image reveals rounded linear ridges (8 -15 km in width and 30 -60 km long. Radar dark material (lava and/or aeolian deposits) fills the region between the ridges. The apparent layered and deformed structures could be oceanic sediments reworked in metamorphic and/or igneous processes after ceasing an aqueous period of Venus' history. Note a canali-type structure in the central part of the image that could be formed by non-silicate magma, such as chloride melt  \n<!-- image -->  \nFig 16.13 Magellan radar image of Tellus Tessera on Venus (NASA/JPL). The arcuate lines of high radar backscatter are marked by gold arrows, and intratessera radar-dark material is marked with blue arrows (Byrne et al. 2021)  \n<!-- image -->  \n<!-- image -->  \nFig 16.14 Radar properties of Maxwell Montes, the highest region on Venus (Treiman et al. 2016). Black symbols are Magellan radar emissivity for all Maxwell Montes from Klose et al. (1992). Red symbols and legend are Magellan SAR radar reflectance coefficients and elevations for the three noodles studied by Treiman et al. (2016). The range of radar backscatter values above ∼ 4 km could reflect the roughness of the topography  \n<!-- image -->  \nFig. 16.15 The potential chemical reactions involving S-bearing compounds in the middle and lower atmosphere and in a permeable surface layer on Venus. The orange rectangle represents clouds, and the grey box designates surface materials. The dotted oval signifies near-surface conditions where gas phase and gas-solid type chemical equilibria might occur. White rectangles denote gases, and ovals illustrate aerosol components. Red arrows indicate photochemical reactions.  Blue and dark yellow arrows represent oxidation and reduction reactions, respectively. Double arrows illustrate thermochemical equilibria between compounds.  Only the significant species and reaction pathways are shown, with details provided in Tables 16.6 and 16.7. H2S, HS, S2O, S-C, and S-Cl-O gases are not depicted. Secondary minerals, such as pyrite, anhydrite, thenardite, magnetite, and hematite, which form via gas-solid reactions, are presented; other secondary phases are discussed in Sect. 16.3.2 . Some pathways, such as fates of OCS and Sn and gas-solid reactions that involve Fe sulfides and oxides, are hypothetical  \n<!-- image -->  \nFig. 16.16 The fate of major S-bearing species in the Hadley-cell type atmospheric circulation on Venus. Global calculation primarily involves the high-density sub-cloud atmosphere but also impacts the clouds and the upper atmosphere. Polar upwellings are suspected (Oschlisniok et al. 2021; Marcq et al. 2023). OCS forms through CO reactions (16.29) and (16.30) with sulfur gases below ~10 km. Low latitude upwelling transports SO2 and OCS towards the clouds. In the clouds, SO2 oxidizes to sulfuric acid (Fig. 16.15; Eqs. (16.7), (16.8)). OCS is significantly consumed in reactions with SO3 in the mid-30s km (Eqs. (16.19), (16.20), (16.21)), leading to CO, SO2, and potentially S2. As a result of these reactions, concentrations of CO, SO2, and possibly S2 increase with latitude in the middle 30 km. High-altitude downwelling delivers CO-, SO2-, and Sn-enriched air to the low sub-cloud atmosphere. In addition to the potential formation of S2 in reaction (16.20), Sn gases form via the thermal dissociation of S8 near the cloud deck (Fig. 16.15). The cycle concludes with OCS formation in a deep equatorward airflow that could approach chemical equilibrium in a thin near-surface layer. While gas transport via eddy diffusion allows for chemical disequilibria that drive reactions throughout the atmosphere (e.g., Krasnopolsky 2007; Bierson and Zhang 2020; Yung et al. 2009), the global circulation results in latitudinal gradients of OCS, CO, and SO2 that may not be explained by eddy diffusion only  \n<!-- image -->  \n<!-- image -->  \nFig. 16.17 The modeled chemical composition of the lower atmosphere of Venus (Krasnopolsky 2013). Mixing ratios of SO2 (130 ppmv) and HCl (0.5 ppmv) are constant and not shown. Sx stands for the sum of Sn (n = 1 -8)  \n<!-- image -->  \n<!-- image -->  \nFig. 16.18 Sulfur-bearing gases in the middle and upper atmosphere of Venus based on models and observations (Krasnopolsky 2012). Observations: (1) Pioneer Venus, Venera 15, HST and rocket data (Esposito et al. 1997); (2) mean results of Sandor et al. (2010); the observed SO2 varies from 0 to 76 ppb and SO from 0 to 31 ppb; (3) Krasnopolsky (2010); (4) Marcq et al. (2011); (5) Belyaev et al. (2012). See Table 16.1 for used observational methods  \n<!-- image -->  \nFig. 16.19 The schematic diagram showing possible formation pathways of condensed sulfur (Sx) in clouds of Venus (Pinto et al. 2021). Parent molecules are SO2 and OCS, and intermediate species are in the circles. Species needed for the reactions are next to the arrows. X can be O, H, NO, S, SO, or S2  \n<!-- image -->  \n<!-- image -->  \nFig. 16.20 The important chemical pathways for S-bearing species in the middle and upper atmosphere of Venus (Zhang et al. 2012)  \n<!-- image -->  \n<!-- image -->  \nFig. 16.21 The estimated steady-state profiles for S-bearing species (Frances-Morrenis et al. 2022).  Pinto et al. (2021) S2 formation mechanism is off, and the reaction of SO with S2O (Eq. (16.48)) becomes the dominant pathway for S2. Solid curves correspond to a higher reaction rate constant that is consistent with the ab initio calculation of Frances-Morrenis et al. (2022) and yields substantial S2  \n<!-- image -->  \n<!-- image -->  \nFig. 16.22 The stability of rock-forming silicates, carbonates and sulfates of Ca and Mg, and solids in the Fe-S-O system as functions of fugacities of SO2 and S2 (bars) at the conditions of Venus' modal radius (740 K, 95.6 bars, 0.6 km below 6052 km level). The equilibrium lines correspond to reactions analogous to (16.71) and (16.74). The dashed lines correspond to (16.71) at the specified activity of anorthite, a (An); activities of other solids are assumed to be unity. The f S2 value for the Py-Pyrr equilibrium (16.76) is from Hong and Fegley (1998). The arrow bars show the uncertainties of selected equilibria due to the thermodynamic data of compounds. The Venus' parallelogram corresponds to measured x SO2 of 100 -200 ppmv (Table 16.1) and f S2 corresponding to the S2-CO2-SO2-CO gas equilibrium at 8 and 17 ppmv CO (Table 16.2). Symbols 7 to 10 represent the corresponding models in Table 16.2. The box around symbol 10 reflects x S3 of 9 -13 pptv (Krasnopolsky 2013), S3-S2 gas equilibrium, and 100 -200 ppmv SO2. The diagram illustrates the instability of Ca pyroxenes and carbonates and the stability of Mg silicates and magnesite concerning sulfatization. The stability field of anorthite is close to Venus' conditions , and plagioclase with an intermediate composition could be in equilibrium with atmospheric SO2 and S2 at the lowlands. Venus' environments are close to conditions at which magnetite and pyrite and/or magnetite and hematite coexist. This match suggests oxidation and pyritization of exposed ferrous silicates, pyrrhotite, and mafic glasses at the atmosphere-surface interface at lowlands (Fig. 16.15)  \n<!-- image -->  \nFig. 16.23 The mineral phase equilibria for the basalt composition as a function of oxygen and sulfur fugacities at conditions of Venus' modal radius (740 K and 95.6 bars) from Semprich et al. (2020). The black box represents a broad range of fugacities of interest at the surface. Hematite -red, magnetite -brown, pyrite -dark yellow, pyrrhotite -bright yellow, ilmenite + magnetite -purple, ilmenite -blue, ilmenite + pyrrhotite -green. Solid solutions: Cpx -clinopyroxene, Opx -orthopyroxene, Ol -olivine, Pl-plagioclase, Kfs -K-feldspar, Po-pyrrhotite, Ilm -ilmenite; non-solid solutions: anh -anhydrite, crd -cordierite, hem -hematite, mag -magnetite, qz -quartz, rt-rutile  \n<!-- image -->  \nFig. 16.24 Stability fields of sulfates and chlorides at the surface of Venus at 740 K and 95.6 bars. Venus atmospheric conditions (red box) correspond to 25 -230 ppmv SO2 (Table 16.1) and 4 ± 0.03 ppmv HCl (Krasnopolsky 2010a). The circle symbol stands for 150 ppmv SO2 and 0.4 ppmv HCl. Phase equilibrium lines correspond to reactions analogs to (16.68) at 30 ppmv H2O and f O2 of 10 -21.36 , corresponding to 0.4 ppmv S2 in equilibrium with 150 ppmv SO2 (Table 16.2, Model 8). The upper and lower ends of the error bars are for f O2 of 10 -21.7 and 10 -20.0 from Fegley et al. (1997b), respectively, and corresponding 2 ppmv and 0.8 ppbv S2. Lower f O2 and higher f S2 enhance the stability of chlorides. The diagram shows the instability of Ca and Mg chlorides concerning sulfurization. Chlorides of K and Na could only be stable if atmospheric x SO2 is below ~70 and ~60 ppmv, respectively  \n<!-- image -->  \n<!-- image -->  \nFig. 16.25 The stability fields of solids in the Fe-S-O system as functions of fugacities of O2 and S2 (bars) at conditions of Venus' modal radius. The equilibrium lines correspond to reactions analogous to Eqs. (16.66) and (16.67). The value of f S2 for the Py-Pyrr equilibrium (16.76) is from Hong and Fegley (1998). Venus' atmospheric conditions are in the caption of Fig. 16.22. If the composition of Venus' near -surface atmosphere corresponds to Models 7 and 8 (Table 16.2), the primary rock's pyrrhotite will oxidize to a magnetite -pyrite assemblage.  \n<!-- image -->  \nPyrrhotite could oxidize to a magnetite-hematite assemblage if the composition corresponds to Models 9 and 10.  The diagram shows that ferric and ferrous sulfates are unstable at the surface. Ferric sulfate grains sunk from the clouds decompose into hematite  \nTemperature, K  \n<!-- image -->  \nFig. 16.26 .  The relative facility of SO3 and SO2 as a function of temperature and f O2 in a steam atmosphere on early Venus. Other curves show conditions of magnetite-hematite and Ni-NiO buffers. The arrow illustrates conditions in which sulfur exists in the sulfate form in silicate magmas that could have formed at the surface because of greenhouse heating. The plot demonstrates that SO3 is more abundant than SO2 only in a high-temperature and O2-rich atmosphere. It f O2 did not exceed values determined by the Mt-Hem buffer (16.66) in surface magmas or altered rocks, SO2 dominated over SO3  \nNiQ  \n<!-- image -->  \nFig. 16.27 Stability fields of S-bearing solutes, pyrite, and native sulfur in seawater 25 °C, 1 bar total pressure, and total S content of 10 -3 and 10 -6 mole/kg H2O (Rickard 2012a). The values in the pyrite stability field correspond to the total Fe content in the system  \n<!-- image -->  \n<!-- image -->  \nFig. 16.28 Speciation of Venus' volcanic gases as a function of C/H ratio at 92.1 bars and 1500 K (Zolotov and Matsui 2002). The elemental composition of gases corresponds to the reconstructed 1918 analysis of Kilauea magma lake gases (Gerlach 1980) with variable C/H ratio. Kilauea emissions are among the most Hdepleted gases. Venus' counterparts could be more H-depleted than Kilauea gases owing to both H deficiency in the interior and suppressed degassing of H2O at the ambient pressure. The oxidized (QFM) and reduced (IW) oxidation states are f O2 endmembers for Venus' mafic melts suggested by the MnO/FeO ratio in the Venera 13, 14 and Vega 2 probes (Schaefer and Fegley 2017, Sect. 16.2.2.1 )  \n<!-- image -->  \nFig. 16.29 S peciation of possible Venus' volcanic gases as a function of S/C ratio. The elemental composition of gases is the same as in Fig. 16.28, and the C/H ratio corresponds to that in the atmosphere. The GL2014 arrow at S/C ratio of 3 × 10 -3 relates to the degassing model of Gaillard and Scaillet (2014) for ~100 bars. As discussed in the text, comparable S and C contents are more likely for Venus, possibly like Kilauea lava lake samples. Higher S/C ratios could represent a major early CO2 release (e.g., from a magma ocean) and its minor volcanic degassing during the geological evolution  \n<!-- image -->  \nlog10  \n<!-- image -->  \nFig. 16.30 Volcanic/Venus atmospheric ratios of S-bearing gases. In a , the volcanic gas composition represents four endmembers in terms of f O2 and H/C atomic ratio from Fig. 16.28. Hrich gases correspond to Kilauea-like compositions, and Hdepleted gases have Venus' atmospheric H/C ratio. The atmospheric gas composition corresponds to Model 8 in Table 16.2. In b , the volcanic gas represents the S-, H-depleted composition shown in Fig. 16.29 at an S/C ratio of 3 × 10 -3 from Gaillard and Scaillet (2014), and the atmospheric composition is for Models 8 and 9 in Table 16.2. The figure demonstrates that volcanic gases are compositionally different from the near-surface air and commonly more abundant than their atmospheric counterparts. The difference is lower in (b) owing to the supposedly suppressed high-pressure degassing of S-bearing species estimated by Gaillard and Scaillet (2014)  \n<!-- image -->  \nFig. 16.31. Possible future changes in concentrations of atmospheric SO2 and the CO/CO2 ratio controlled by the magnetite-pyrite-hematite equilibrium (e.g., (16.83), (16.84)) in surface materials at 95.6 bar. The fugacity of SO2 in the current near-surface atmosphere reflects a range of measured SO2 concentrations (Table 16.1) with uncertainties. The triangle symbol shows the nominal conditions of the mineral equilibrium at 740 K (see Table 16.11 for other gases). The double arrow shows the potential f SO2 error bar due to the uncertainty of the thermodynamic data of minerals in (16.84). The circle symbol corresponds to the partial pressure of SO2 of 1 bar at 786 K. The dotted curve shows possible conditions through the thermal decomposition of pyrite in a permeable surface layer. Present atmospheric data on SO2 data are from Oyama et al. (1980) and Gel'man et al. (1980). The figure is modified after Gillmann et al. (2022)  \n<!-- image -->  \n<!-- image -->"}