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Sulfur Chemistry in the Middle Atmosphere of Venus
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Xi Zhang 1*, Mao Chang Liang 2,3,4, Frank P. Mills 5, Denis A. Belyaev 6,7 and Yuk L.
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Yung 1
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Division of Geological and Planetary Sciences, California Institute of Technology,
Pasadena, CA, 91125, USA
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Research Center for Environmental Changes, Academia Sinica, Taipei, Taiwan
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Graduate Institute of Astronomy, National Central University, Zhongli, Taiwan
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Institute of Astronomy and Astrophysics, Academia Sinica, Taipei, Taiwan
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RSPE/AMPL, Mills Rd., Australian National University, Canberra, ACT 0200, Australia
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LATMOS, CNRS/INSU/IPSL, Quartier des Garennes, 11 bd. d’Alembert, 78280,
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Guyancourt, France
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Space Research Institute (IKI), 84/32 Profsoyuznaya Str., 117997, Moscow, Russia
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*Corresponding author: xiz@gps.caltech.edu
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Abstract
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Venus Express measurements of the vertical profiles of SO and SO2 in the middle
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atmosphere of Venus provide an opportunity to revisit the sulfur chemistry above the
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cloud tops (~58 km). A one dimensional photochemistry-transport model is used to
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simulate the behavior of the whole chemical system including oxygen-, hydrogen-,
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chlorine-, sulfur-, and nitrogen-bearing species. A sulfur source is required to explain the
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SO2 inversion layer above 80 km. The evaporation of the aerosols composed of sulfuric
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acid (model A) or polysulfur (model B) above 90 km could provide the sulfur source.
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Measurements of SO3 and SO (a1∆→X3∑) emission at 1.7 μm may be the key to
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distinguish between the two models.
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Keywords: Venus, atmosphere; Atmospheres, chemistry; Photochemistry; Atmospheres,
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composition;
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1: Introduction
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Venus is a natural laboratory of sulfur chemistry. Due to the difficulty of observing the
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lower atmosphere, we are still far from unveiling the chemistry at lower altitudes (Mills
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et al., 2007). On the other hand, the relative abundance of data above the cloud tops (~60
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km) allows us to test the sulfur chemistry. Mills et al. (2007) summarized the important
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observations before Venus Express and gave an extensive review of the chemistry of
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sulfur on Venus.
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Recently, measurements of Venus Express and ground-based observations have greatly
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improved our knowledge of the sulfur chemistry. Marcq et al. (2005; 2006; 2008)
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reported the latitudinal distributions of CO, OCS, SO2 and H2O in the 30-40 km region.
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The anti-correlation of latitudinal profiles of CO and OCS implies the conversion of OCS
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to CO in the lower atmosphere (Yung et al., 2009). Using the latitudinal and vertical
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temperature distribution above 70 km obtained by Pätzold et al. (2007), Picciani et al.
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(2008) deduced the dynamic structure, which is consistent with a weak zonal wind field
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above ~70 km. The discovery of the nightside warm layer (Bertaux et al., 2007) is strong
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evidence of substantial heating in the upper mesosphere. Near the antisolar point, this
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heating is consistent with the existence of a subsolar-antisolar (SSAS) circulation in the
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lower thermosphere (above 90 km). However, the nightside warm layer has been reported
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in SPICAV observations at all observed local times and latitudes and does not appear to
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be consistent with ground-based submillimeter observations (Clancy et al., 2008).
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Through the occultation technique, Solar Occultation at Infrared (SOIR) and the
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Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV)
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onboard Venus Express carry out measurements of the vertical profiles of the major
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species above 70 km, including H2O, HDO, HCl and HF (Bertaux et al., 2007), and CO
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(Vandaele et al., 2008), SO and SO2 (Belyaev et al., 2008; 2010). Aerosols are found to
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be in a bimodal distribution above 70 km (Wilquet et al., 2010). These high vertical
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resolution profiles are obtained mainly in the polar region. The ground-based
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measurements also provide valuable information. Krasnolposky (2010a; 2010b) obtained
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spatially resolved distributions of CO2, CO, HDO, HCl, HF, OCS and SO2 at the cloud
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tops from the CSHELL spectrograph at NASA/IRTF. Sandor et al. (2010) reported
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ground-based submillimeter observations of SO and SO2 inversion layers in the lower
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thermosphere (above 85 km). The submillimeter results are qualitatively consistent with
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the vertical profiles from Venus Express (Belyaev et al., 2010). However, only the
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smallest SO and SO2 abundances inferred from the SPICAV observations (Belyaev et al.,
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2010) are quantitatively similar to those inferred from the submillimeter observations
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(Sandor et al., 2010). Spatial and temporal variability may explain at least some of the
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differences among the observations (Sandor et al., 2010), but detailed inter-comparisons
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are required. These measurements and the proposed correlation of upper mesosphere SO2
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abundances with temperature open up new opportunities to study the photochemical and
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transport processes in the middle atmosphere of Venus.
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Sandor et al. (2010) found that SO and SO2 inversion layers cannot be reproduced by the
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previous photochemical model (Yung and DeMore, 1982). Therefore they suggested that
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the photolysis of sulfuric acid aerosol might directly produce the gas phase SOx. A
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detailed photochemical simulation by Zhang et al. (2010) shows that the evaporation of
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the H2SO4 aerosol with subsequent photolysis of H2SO4 vapor could provide the major
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sulfur source in the lower thermosphere if the rate of photolysis of H2SO4 vapor is
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sufficiently high. Their models also predict supersaturation of H2SO4 vapor pressure
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around 100 km. The latest SO and SO2 profiles retrieved from the Venus Express
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measurements agree with their model results (Belyaev et al., 2010). This mechanism
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reveals the close connection between the gaseous sulfur chemistry and aerosols.
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Previously sulfuric acid was considered as the ultimate sink of gaseous sulfur species in
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the middle atmosphere. If it could also be a source, the thermodynamics and
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microphysical properties of H2SO4 must be examined carefully. Alternatively, if the
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polysulfur (Sx) is indeed a significant component of the aerosols as the unknown UV
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absorber, Carlson et al. (2010) estimated that the elemental sulfur is about 1% of the
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H2SO4 abundance. This might be also enough to produce the sulfur species if there is a
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steady supply of Sx aerosols to the upper atmosphere.
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The purpose of this paper is to use photochemical models to investigate the sulfur
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chemistry in the middle atmosphere and its relation with aerosols based on our current
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knowledge of observational evidences and laboratory measurements. We will introduce
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our model in section 2. In section 3, we will provide a detailed discussion on the two
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possible sulfur sources above 90 km, H2SO4 and Sx, respectively, the roles they play in
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the sulfur chemistry, their implications and how to distinguish the two sources by the
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future observations. The last section provides a summary of the paper and conclusions.
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2: Model Description
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Our photochemistry-transport model is based on the one dimensional Caltech/JPL
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kinetics code for Venus (Yung and DeMore, 1982; Mills, 1998) with updated chemical
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reactions. The model solves the coupled continuity equations with chemical kinetics and
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diffusion processes, as functions of time and altitude from 58 to 112 km. The atmosphere
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is assumed to be in hydrostatic equilibrium. We use 32 altitude grids with increments of
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0.4 km from 58 to 60 km and 2 km from 60 to 112 km. The diurnally averaged radiation
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field from 100-800 nm is calculated using a modified radiative transfer scheme including
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the gas absorption, Rayleigh scattering by molecules and Mie scattering by aerosols with
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wavelength-dependent optical properties (see Appendix A). The unknown UV absorber is
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approximated by changing the single scattering albedo of the mode 1 aerosol beyond 310
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nm, as suggested by Crisp (1986). Because the SO, SO2 and aerosol profiles from Venus
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Express are observed in the polar region during the solar minimum period (2007-2008),
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our calculations are set at a circumpolar latitude (70N) and we use the combination of
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the low solar activity solar spectra for the time period of the Spacelab 3 ATMOS
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experiment with an overlay of Lyman alpha as measured by the Solar Mesospheric
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Explorer (SME).
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In this study we select 51 species, namely, O, O(1D), O2, O2(1∆), O3, H, H2, OH, HO2,
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H2O, H2O2, N2, Cl, Cl2, ClO, HCl, HOCl, ClCO, COCl2, ClC(O)OO, CO, CO2 , S, S2 , S3,
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S4, S5, S7, S8, SO, (SO)2, SO2, SO3, S2O, HSO3, H2SO4, OCS, OSCl, ClSO2, four
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chlorosulfanes (ClS, ClS2, Cl2S, and Cl2S2) and eight nitrogen-containing species (N, NO,
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NO2, NO3, N2O, HNO, HNO2, and HNO3). The chlorosulfanes (SmCln) are included
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because they open an important pathway to form S2 and polysulfur Sx(x=2→8) in the region
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60-70 km (Mills and Allen, 2007), although there are significant uncertainties in their
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chemistry. The chlorosulfane chemistry is not important for the sulfur cycles above ~80
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km because the SmCln abundances are low. Nitrogen-containing species, especially NO
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and NO2, can act as catalysts for converting SO to SO2 and O to O2 in the 70-80 km
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region (Krasnopolsky, 2006). The nitrogen chemistry is not important above 80 km.
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In Zhang et al. (2010), the chemistry is simplified because (SO)2, S2O and HSO3 are
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considered only as the sinks of the sulfur species. This may not be accurate enough for
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the lower region chemistry below 80 km where there seems to be difficulty in matching
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the SO2 observations below 80 km. Instead, a full set of 41 photodissociation reactions is
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used, along with about 300 neutral chemical reactions, as listed in Table 1 and 2
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respectively. We take the ClCO thermal equilibrium constant from the 1-sigma model in
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Mills et al. (2007) so that we can constrain the total O2 column abundances to around
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2×1018 cm-2. In addition, we introduce the heterogeneous nucleation processes of
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elemental sulfurs (S, S2 and polysulfur) because these sulfur species can readily stick to
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sulfuric acid droplets and may provide the source material for the unknown UV absorbers
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(Carlson et al. 2010). But we neglect all the heterogeneous reactions within the
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condensed elemental sulfur species on the droplet surface. The calculation of the
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heterogeneous condensation rates is described in Appendix B. The accommodation
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coefficient  is varied from 0.01 to 1 for the sensitivity study (section 4).
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The temperature profiles are shown in Fig. 1 (left panel). The daytime temperature profile
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below 100 km (solid line) is obtained from the VeRa observations of Venus Express near
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the polar region (71N, figure 1 of Pätzold et al., 2007). Above 100 km the temperature is
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from Seiff (1983). The nighttime temperature profile (dashed line) above 90 km is
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measured by Venus Express in orbit 104 at latitude 4° S and local time 23:20 h (black
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curve in the figure 1 of Bertaux et al., 2007). The nighttime temperature profile is used
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only to calculate the H2SO4 saturation vapor pressure.
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[Figure 1] (see page 43)
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In the 1-D model, species are assumed to be mixed vertically by turbulent eddies.
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Molecular diffusion becomes important above the homopause region (~130-135 km). The
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eddy diffusivity increases with altitude in the Venus mesosphere for reasons discussed
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later. Since the zonal wind is observed to decrease with altitude above 70 km due to the
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temperature gradient from pole to equator (Piccialli et al., 2008), we assume that vertical
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transport of gas species is predominantly due to mixing caused by transient internal
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gravity waves, as proposed by Prinn (1975). In this case the eddy diffusion coefficient
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will be inversely proportional to the square root of the number density (Lindzen, 1971).
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Von Zahn et al. (1980) derived the eddy diffusion coefficient ~1.4×1013 [M]-0.5 cm2 s-1
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from of the number density around the homopause region, based on mass spectrometer
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measurements from Pioneer Venus. In our model we use slightly larger values Kz =
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2.0×1013 [M]-0.5 cm2 s-1 above 80 km to include an additional contribution above 80 km
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due to the SSAS circulation. In the region below 80 km. Woo and Ishmaru (1981)
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estimated the diffusivity  4.0×104 cm2 s-1 at ~60 km from radio signal scintillations.
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Therefore our eddy diffusivity profile from 58-80 km is estimated by linear interpolation
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between the values at 80 and 60 km. The eddy diffusion profile is shown in Fig. 1 (right
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panel). However, the vertical advection due to the SSAS circulation might be dominant
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above 90 km. The SPICAV measurements show that the SO2 mixing ratio from ~90-100
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km is almost constant with altitude, which implies a very efficient transport process.
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Since Zhang et al. (2010) has already shown that the model without additional sulfur
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sources above 90 km cannot explain the observed SO2 inversion layer, we focus on two
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models in this study: model A with enhanced H2SO4 abundance above 90 km and model
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B with enhanced S8 abundance above 90 km. In both models we fix the vertical profiles
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of N2, H2O, and H2SO4. The N2 profile is given by a constant mixing ratio of 3.5%. The
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H2O profile (see Fig. 3) is prescribed on the basis of the Venus Express observations
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(Bertaux et al., 2007) above 70 km and is assumed to be constant below. The
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accommodation coefficient of the sulfur nucleation is set to unity (the upper limit) in the
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standard model.
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The H2SO4 saturation vapor pressure (SVP) and the photolysis cross sections are
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discussed in the Appendices C and D, respectively. Model A uses the nighttime H2SO4
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vapor abundance (Fig. A6) and the H2SO4 photolysis cross sections with high UV cross
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section in the interval 195-330 nm (dashed line in Fig. A7). The H2SO4 abundance above
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90 km is scaled to reproduce the observed SO and SO2 data. The day and nighttime Sx
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saturated mixing ratio profiles based on Lyons (2008) are shown in Fig. A6. The S8
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mixing ratio under nighttime temperature could achieve ppb levels at ~96 km. Based on
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this, we fix the S8 profile in model B as the sulfur source. The S8 profile is composed of
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two parts: (1) below 90 km, we use the output from model A, which shows the S8
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concentration is undersaturated; (2) above 90 km, we scale the S8 saturated vapor
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abundances under nighttime temperature to match the observations. We also adopt the
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daytime H2SO4 vapor abundance and the H2SO4 photolysis cross sections with low UV
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cross section in the interval 195-330 nm (solid line in Fig. A7); therefore the H2SO4
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photolysis has almost no contribution to the sulfur enhancement in model B.
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The SO2 and OCS mixing ratios at 58 km are set to 3.5 ppm and 1.5 ppm, respectively.
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The upper and lower boundary conditions for the important species are listed in Table 3.
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We set the HCl as 0.4 ppm at 58 km, which is about factor of 2 larger than the Venus
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Express observations (Bertaux et al., 2007). However, other observations support the 0.4
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ppm HCl (Connes et al., 1967; Krasnopolsky, 2010a). We argue that since ClC(O)OO is
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the key species to convert CO and O2 to CO2 (Mills, et al., 2007), 0.4 ppm HCl is needed
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in our model to constrain the total column abundances of O2.
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3: Model Results
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Enhanced H2SO4 case (Model A)
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In model A, the saturation ratio of H2SO4 is 0.25 (sub-saturated H2SO4 under nighttime
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temperature), corresponding to the peak value ~0.2 ppm at 96 km. Figures 2-8 show the
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volume mixing ratios of oxygen species, hydrogen species (including HOx), chlorine
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species, chlorine-sulfur species, elemental sulfurs, nitrogen species, and sulfur oxides,
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respectively. The observations of SO, SO2 and OCS are also plotted in Fig. 8.
[Figures 2-8] (see pages 44-50)
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The model outputs of oxygen, hydrogen and chlorine species generally agree with the
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previous studies (model C in Yung and DeMore, 1982; the one-sigma model in Mills,
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1998). The differences arise mainly from the differences of the temperature profiles and
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radiation field between the polar region in our model and the mid-latitude in the previous
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models. Compared with the temperature profile from Seiff (1983) used in the previous
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models, the current profile at 70ºN is colder in the upper cloud layer (58-70 km) but
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warmer between 70 km and 80 km. The O2 profile is consistent with the one-sigma model
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in Mills (1998). The concentrations of chlorine species in model A are lower than those
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in model C of Yung and DeMore (1982) mainly because of the lower boundary values of
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HCl (0.8 ppm in their model C and 0.4 ppm in our model). The chlorine-sulfur species in
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our model are slightly larger than those from Mills (1998) near the lower boundary
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because there is more OCS at the lower boundary of the model A. The nitrogen species in
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our model is less than that in model B of Yung and DeMore (1982) simply because their
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lower boundary value of NO is 30 ppb, which is about 6 times larger than the observed
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values, 5.5 ppb on average at 65 km (Krasnopolsky, 2006). Since the polysulfur
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chemistry has been updated based on Moses et al. (2002), model A results are different
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from those of Mills (1998). But the polysulfur chemistry has large uncertainties and more
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laboratory measurements of the reaction coefficients are needed (see discussions in
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section 4). Concentrations of sulfur oxides are obviously different from those of early
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models when we include the sulfur source above 90 km.
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The photolysis of the parent species CO2, OCS, SO2, H2O and HCl, which are transported
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by eddy diffusion from 58 km, provides the sources for the other species. Although the
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sulfur chemistry is closely coupled with the oxygen and chlorine chemistry in the region
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60-70 km, the sulfur species have little effect on the abundances of the oxygen species
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(including HOx) and chlorine species above 70 km, but not vice versa. In other words, the
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model without sulfur chemistry would produce roughly the same amount of oxygen and
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chlorine species as the model with sulfur species does above 70 km where the sulfur
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species are less abundant and can no longer tie up the chlorine and oxygen species. The
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free oxygen and chlorine bearing radicals, such as O, OH, Cl and ClO, are the key
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catalysts in recycling sulfur species. On the other hand, the sulfur species do not act as
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catalysts. Therefore, to some extent the sulfur cycle in the mesosphere can be separated
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from the other cycles above the cloud tops. Fig. 9 illustrates the important pathways of
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the sulfur cycle. For simplicity, the chlorosulfane chemistry and polysulfur chemistry are
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not shown here. See Mills and Allen (2007) and Yung et al. (2009) for detailed
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discussions. A fast cycle, including the photodissociation and oxidization processes,
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exists among the sulfur species. Below 90 km, H2SO4 and Sx act as the ultimate sinks
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rather than the sources of the sulfur species in the region. The total production rate of
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H2SO4 from 58-112 km is 5.6×1011 cm-2 s-1 with a peak value of 1.3×106 cm-3 s-1 around
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64 km, while the total loss rate of gaseous elemental sulfur to aerosol through
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heterogeneous nucleation processes is 2.6×1012 cm-2 s-1, when summed across all Sx that
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are lost, equivalent to a sulfur atom loss rate ~3.3×1012 cm-2 s-1. Therefore, the major sink
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of sulfur species is the nucleation of the polysulfur aerosol. The polysulfur sink decreases
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with altitude mainly because both the aerosol and gaseous sulfur species are less
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abundant at higher altitude (Fig. 6, also see the discussion in Section 4). For reference,
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the Krasnopolsky and Pollack model [1994] requires the H2SO4 production rate of
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2.2×1012 cm-2 s-1. Some previous models range from 9×1011-1×1013 cm-2 s-1 (Yung and
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DeMore, 1982; Krasnopolsky and Pollack, 1994). So the H2SO4 production rate in model
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A is lower than those previously modeled values.
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[Figure 9] (see page 51)
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Below ~65 km, ClSO2 is roughly in chemical equilibrium with SO2. SO2 reacts with the
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chlorine radical: Cl + SO2 + CO2 → ClSO2 + CO2, and ClSO2 reacts with O, S, S2, SO,
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ClSO2, etc. (Reactions R293-R300 in table 2) to return back SO2 and produce chlorine
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species including chlorosulfanes. Above ~65 km, photolysis of SO2 becomes the
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dominant sink, with a minor branch of oxidization to SO3. The three-body reaction O +
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SO produces more SO2, and ClO and ClC(O)OO reacting with SO also play important
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roles in SO2 production. The major production and loss pathways for SO, SO2 and SO3 in
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model A are plotted in Fig. 10.
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[Figure 10] (see page 52)
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The SO and SO2 inversion layers above 80 km are successfully reproduced in model A
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(Fig. 8) although the agreement between the model and observations is not perfect. We
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attribute the discrepancy to the constant H2SO4 saturation ratio above 90 km. The SO2
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measurements imply that the H2SO4 abundance in model A might be underestimated
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above 100 km. The chemistry is mainly driven by the photolysis and reactions with O and
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O2. The fast recycling between the sulfur species can be seen from the largest production
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and loss rates in Fig. 10. At 96 km where the peaks of the reaction rates are, the SO3
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photolysis rate (~1.5×103 cm-3 s-1) and SO3 + H2O rate (~1.5×103 cm-3 s-1) are roughly
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comparable, which implies about half of the sulfur in H2SO4 goes into SOx and produces
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the inversion layers of SO2 and SO. Model A predicts the existence of an SO3 inversion
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layer, with a peak value of ~13 ppb at 96 km. More discussion will be provided in
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Section 4.
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Enhanced Sx case (model B)
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The Sx aerosol might be another possible sulfur source in the upper region because S x
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could react with atomic oxygen to produce SO. But this possibility is more ambiguous
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because: (1) The Sx aerosol has not been indentified although it is the most likely UV
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absorber (Carlson et al. 2010); (2) As shown in the Fig. 6, the production of Sx is mainly
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confined to the region below 65 km. So the Sx in the haze layer might be not sufficient to
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supply the sulfur source; (3) The Sx chemistry has large uncertainty due to the lack of
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laboratory experiments. The reaction coefficients for Sx + O in our model were estimated
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by Moses et al. (2002) based on that for S2 + O.
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The required saturation ratio of S8 is only 3×10-4 in order to produce the SO and SO2
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inversion layer. That means we need only 0.1 ppt Sx vapor above 90 km. The blue lines in
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Fig. 11 show that model B could also reproduce the inversion layers above 80 km. The
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SO and SO2 profiles (and other sulfur species) from models A and B are really similar
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mainly due to the fast sulfur cycle. However, there is a large difference in SO3 profiles
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because the SO3 in model B is derived mainly from SO2 but not from the photolysis of
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H2SO4 as in model A. The SO3 at 96 km predicted by model B is ~0.1 ppb, which is two
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orders of magnitude less than that in model A. Therefore, future measurement of SO3
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could distinguish the two mechanisms.
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[Figure 11] (see page 53)
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The major production and loss pathways for SO, SO2 and SO3 in model B are plotted in
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Fig. 12. Note that not only S8 + O produces SO but other Sx derived from S8 also react
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with O to produce SO, so in fact one S8 gas molecule could eventually produce about
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eight SO molecules. The major differences between models A and B are the SO and SO3
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production mechanisms.
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[Figure 12] (see page 54)
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4: Discussion
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Summary of Chemistry above 80 km
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The results of models A and B are summarized in Table 4.
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For model A, the simplified SOx chemistry from Fig. 9 can be illustrated as:
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
à ààEvaporation
àà àà àà ààÜ
à ààhà
àà SO3 á
à ààhàà
Ü
à ààhàà
Ü
à ààhààÜ
à àà àà àà ààÜ
Aerosol á
à H 2 SO4 á
àÜ
à SO2 á
à SO á
à Sá
à Sx
Condensation
H O
O
O
O
O(weak )
2
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2
Similar to model A, the chemistry in model B:
H2O
O
à ààEvaporation
ààààààààÜ
à ààààOà
à ààOàà
Ü
à ààOàà
Ü
à àààà
ààààÜ
Aerosol á
à Sx á
àààààÜ
à SO á
à SO2 á
à SO3 á
à H 2 SO4
2Ü
Condensation
h
h
h (weak )
ààà
S á à àà SO
h
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In fact OCS, S2O and (SO)2 are also in photochemical equilibrium with the species above
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but not shown here (see Fig. 9). Therefore, the fast cycle allows us to derive the ratios of
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the sulfur species above 90 km analytically by equating the production and loss rates for
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each species (the subscripts refer to the reaction numbers in Tables 1 and 2):
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[OCS] k183[ClCO]  k185 [CO][M ]

S
k302 [S]  J 31
(1)
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[S]
J 27

[SO] k166 [O2 ]
(2)
12
[SO]
J 29

[SO2 ] k240 [O][M ]
(3)
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[SO2 ] J 30  k177 [H 2O]

[SO3 ]
k258 [O][M ]
(4)
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But in model A, SO3 is approximately independent of other SOx because of the photolysis
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of sulfuric acid is the principal source:
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In model B:
[SO3 ] 
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J 32 [H 2 SO4 ]
J 30  k177 [H 2O]
(5)
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Again, SO3 is the key species to distinguish the two pathways because other sulfur
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species are closely coupled to SO2 no matter what causes the inversion layers.
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The S2O and (SO)2 chemistry is less clear so it needs a more careful study in the future.
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In models A and B, the steady-state results are given by:
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[S2O]
k250 [O]

[(SO)2 ] J 33  k266 [O]
(6)
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[(SO)2 ]
k249 [SO][M ]

[SO]
k246 [O]  k254 [M ]
(7)
382
The total column abundance of O2 is about ~2.2×1018 cm-2 in models A and B. The
383
atomic oxygen (O) column production rate above 80 km is ~4.3×1012 cm-2 s-1. The O flux
384
required to reproduce the mean O2 emission of 0.52 MR in the nightside is ~2.9×1012 cm-
385
2 -1
386
latitude) instead, we obtain ~6×1012 cm-2 s-1. This suggests that about 50% of the O atoms
387
produced in the dayside are transported to the nightside and recombine to form O2.
388
Therefore, the transport process is at least as fast as the chemical loss processes. The loss
389
timescale of O is ~105-106 s above 80 km. The transport timescale is estimated to be ~104
390
s, based on the SSAS downward velocity ~ 43 cm/s around 100 km in the night side from
391
Bertaux et al. (2007) and the scale height of 4 km in the lower thermosphere.
s (Krasnopolsky, 2010c). Note that our calculation is at 70N. If we use 45N (mid-
392
393
Both the SO and SO2 profiles are derived from the observations above 90 km. They
394
provide a test of the three-body reaction SO + O + M → SO2 + M in the low temperature
13
395
region, for which there are no laboratory measurements. The reaction coefficient k240,0
396
derived from observations is about 3×10-30 at 100 km (168 K). The value adopted in our
397
models is from Singleton and Cvetanovic (1988) at 298 K, k240,0 = 4.2×10-31 cm6 s-1and
398
k240,∞ = 5.3×10-11 cm3 s-1, and has been enhanced by a factor of 8.2 for the third-body
399
CO2. This reaction coefficient produces the [SO2]/[SO] ~2, which lies in the Venus
400
Express observations range (Beyleav et al., 2010) and the dayside submillimeter
401
observation range (Sandor et al., 2010). The nightside submillimeter observations,
402
however, find [SO2]/[SO] ~15-50 (Sandor et al., 2010). It implies that this reaction may
403
have no or very weak temperature dependence between 160-300 K. Grillo et al. (1979)
404
and Lu et al. (2003) measured the temperature dependence in the high temperature region
405
(~300-3000 K) and the two papers provided the dependence as T-1.84 and T-2.17,
406
respectively. However, this temperature dependence is too steep for the low temperature
407
region.
408
409
One puzzle from the observed [SO2]/[SO] ratio is that the ratio seems to increase with
410
temperature at 100 km (Belyaev, et al., 2010), although this difference is within the
411
uncertainty associated with fewer measurements in the high temperature region (the T2
412
(180-185 K) and T3 (190-192 K) regions). In addition, there is difficulty in separating the
413
SO signal from the SO2 signal in the spectrum.
414
415
Sensitivity Study
416
417
Due to the uncertainty in the adopted value of the accommodation coefficient, we test the
418
sensitivity of the model by slowing down the heterogeneous nucleation processes
419
(reducing ). As  decreases, the formation of Sx aerosols is slower so there are more
420
sulfur species in the gas phase. Therefore, the model requires less SO2 at the lower
421
boundary to reproduce the SO2 observations. The left panel of Fig. 13 summarized the
422
cases with 1 (model A), 0.1 and 0.01 . All three cases show that the major sink of
423
sulfur species is the polysulfur aerosol. As the polysulfur nucleation process slows down,
424
the H2SO4 production rate also decreases because of the lower SO2 abundances around
425
62-64 km where the H2SO4 production peak is. One interesting phenomenon is that, when
14
426
the nucleation process is slower, the production rate of condensed Sx aerosol is less but
427
the equivalent sulfur atom loss rate increases. This is because more high order polysulfur
428
molecules are produced and condense rapidly to form aerosols. Therefore, it is the result
429
of the competition between the cycling of polysulfur and the nucleation processes. The
430
region above 85 km is almost unaffected because the polysulfur sink at those levels is
431
negligible.
432
[Figure 13] (see page 55)
433
434
The middle and right panels of Fig. 13 show the sensitivity to different lower boundary
435
values of SO2 (2.5-5 ppm) and OCS (0.3-5 ppm). These cases all lie in the observational
436
error bars. The change of SO2 boundary conditions mainly impacts the region below 85
437
km. The OCS boundary conditions affect only the upper cloud region.
438
439
The major uncertainties of model A arise from the H2SO4 vapor abundances and the
440
photolysis rate of H2SO4. Since half of the sulfur in H2SO4 goes into SOx, we would
441
expect a linear relationship between [SO2] and the product of J32 and [H2SO4]. In model
442
A, the J32 and [H2SO4] at 96 km are 6.7×10-6 s-1 and 5.0×108 cm-3, respectively, so
443
J32[H2SO4] is ~3.3×103 cm-3 s-1. For model B, although the relationship between the S8
444
and SO2 abundances is not derived explicitly, we also expect a linear relationship
445
between the two species because all the major sulfur oxides in model B above 90 km are
446
linearly dependent with each other. The reaction coefficients of S8 + O (k233) and [S8] at
447
96 km are 7.0×10-12 cm3 s-1 and 7.1×102 cm-3 respectively, so k233[S8] (the product of k233
448
and [S8]) is ~5.0×10-9 s-1. Assuming that all eight sulfur atoms in S8 eventually go into
449
SO, the production rate is ~2×103 cm-3 s-1 given the O abundance is ~4×1010 cm-3 at 96
450
km. This value is of the same order of magnitude as J32[H2SO4] from model A. Therefore,
451
both models suggest that this magnitude of sulfur source is necessary to explain the
452
inversion layers.
453
454
Sulfur Budget above 90 km
455
15
456
The recycling of aerosols from the region below 90 km is essential to maintain the
457
inversion in steady state because sulfur will diffuse downward due to the inverted mixing
458
ratio gradient. The production rates of H2SO4 and Sx aerosol for models A and B are
459
shown in Fig. 14. The H2SO4 production rate in model A has a secondary peak at 96 km
460
where the SO3 peak is located. But the Sx aerosol production rate above 90 km is small
461
because the nucleation process is really slow when there are few due aerosol particles to
462
serve as condensation nuclei.
463
[Figure 14] (see page 56)
464
465
In model A, the net column loss rate of H2SO4 vapor above 90 km is ~5×108 cm-2 s-1,
466
corresponding to ~50% of the column photolysis rate of H2SO4 in that region. Only ~2%
467
sulfur is converted into polysulfur aerosol. So the total downward sulfur flux is ~5×108 to
468
maintain a steady-state above 80 km. For reference, the 5×108 cm-2 s-1 H2SO4 loss rate
469
above 90 km is roughly equal the H2SO4 column production rate in the region of 78-90
470
km through the hydration of SO3. However, it is difficult to transport the H2SO4 vapor
471
upward from below 90 km to compensate the loss above because the vapors will quickly
472
condense into the aerosols. Instead, the aerosols must be transported upward on the
473
dayside by the SSAS circulation. Assuming that all the aerosols above 90 km are the
474
mode 1 aerosols with mean radius ~0.2 m and density 2 g cm-3, the aerosol column loss
475
rate is ~1 cm-2 s-1. The column abundance of mode 1 aerosol above 90 km from Wilquet
476
et al. (2010) is ~5.0×106 cm-2. Therefore the aerosol lifetime is ~100 earth days in model
477
A. The loss rate implies that the upward aerosol flux is also ~1 cm-2 s-1 across 90 km to
478
supply the aerosol budget. Provided that the concentration of mode 1 aerosol at 90 km is
479
~10 cm-3, the estimated flux is equivalent to an effective upward transport velocity ~0.1
480
cm s-1. If the SSAS circulation dominates the upper atmosphere, this velocity is readily to
481
be achieved. The SSAS circulation with a vertical velocity of ~43 cm s-1 (Bertaux et al,
482
2007) might be very efficient for transporting the aerosols in the upper regions.
483
484
In model B, the total equivalent sulfur column loss rate of the Sx vapor above 90 km is
485
~4×108 cm-2 s-1, which is roughly the same magnitude as the net sulfur flux from H2SO4
486
photolysis in the model A, but the production rate of Sx aerosol is only ~1×107 cm-2 s-1.
16
487
The H2SO4 aerosol production rate in model B is ~2×107 cm-2 s-1. Therefore, most of the
488
sulfur from Sx aerosol diffuses downward as SO2 and SO. For reference, the equivalent
489
sulfur column production rate of the Sx aerosol above 78 km is ~4.0×108 cm-2 s-1.
490
Therefore, if those aerosols can be transported upward, it would be enough to compensate
491
for the loss in the region above 90 km. Assuming that all the polysulfur aerosol above 90
492
km has a mean radius of ~0.1 m (half of the H2SO4 aerosol) and the density is 2 g cm-3,
493
the aerosol column loss rate is ~3 cm-2 s-1. If the polysulfur aerosol abundance is about
494
1% of that of H2SO4 aerosol, Sx aerosol above 90 km will be removed in ~1 earth day.
495
The estimated upward flux cross the 90 km level is equivalent to an effective upward
496
transport velocity ~30 cm s-1, which is roughly in the same magnitude of the downward
497
velocity at 100 km in the nightside from Bertaux et al. (2007). Liang et al. (2009) derived
498
the velocity ~1 cm s-1 at 100 km, which might be too small.
499
500
Timescales
501
502
The dynamics in the 1-D photochemistry-transport model is only a simple
503
parameterization for the complicated transition zone between 90-100 km. The aerosol
504
microphysics is also simplified because we just assume the instantaneous condensations
505
of H2O and H2SO4 and ignore the aerosol growth and loss processes. Future 2-D models
506
including SSAS, zonal wind transport, microphysical processes and photochemical
507
processes for both the dayside and nightside hemisphere might be sufficient to represent
508
all the dynamical and chemical processes in the upper regions. We estimate some typical
509
timescales here.
510
511
(1) Transport: The timescale for the SSAS transport SSAS is ~104 s (section 3) for vertical
512
transport over the 4 km scale height near 100 km. Zonal transport timescale due to RZ
513
flow RZ is ~105 s for transport from the subsolar point to the antisolar point, provided
514
that the thermal wind velocity is ~50 m s-1 by Piccialli et al. (2008) based on the
515
cyclostrophic approximation. Eddy diffusion timescale eddy is also ~105 s near 100 km
516
(Fig. A3).
517
17
518
(2) Aerosol condensation: We assume the condensation is dominated by the
519
heterogeneous nucleation with timescale (Fig. A4) cond~104-105 s for the accommodation
520
coefficient =1. cond is inversely proportional to  in the free molecular regime for the
521
upper region. Therefore, the lower value of  might be more appropriate for the H2SO4
522
condensation in model A since the mechanism assumes that the nighttime H2SO4 vapor
523
could be transported to the dayside and undergo photolysis. Homogeneous nucleation
524
may be important in the dayside since both of H2SO4 and Sx are highly supersaturated.
525
For example, the dayside saturation ratios at 100 km are about 106 and 104 for H2SO4 in
526
model A and S8 in model B, respectively. In model A, these dayside condensation
527
processes will compete with the photodissociation of H2SO4. However, the homogeneous
528
nucleation rate markedly depends on the SVP, which is a steep function of temperature
529
but not well determined in the lower temperature range. If the dayside condensation
530
process were fast, a zonal gradient of the H2SO4 vapor abundance from the nightside to
531
the dayside would be expected.
532
533
(3) Chemistry: H2SO4 photolysis timescale photo depends on the cross section. For model
534
A, photo is ~105 s. Sx + O timescale sx+o depends on the reaction coefficient and the
535
atomic oxygen abundance, Sx+O is ~1-10 s in the model B. That is why ~0.1-1 ppt Sx
536
could provide a sulfur source as large as the photolysis of ~0.1 ppm H2SO4 does. The
537
reaction between polysulfur and atomic oxygen is so fast that it has to happen in the
538
nightside. However, as shown above, whether the circulation could support the Sx aerosol
539
upward flux across 90 km needs more future studies.
540
541
OCS above the cloud tops
542
543
The OCS mixing ratio in the upper cloud layer is puzzling. OCS originates from the
544
lower atmosphere. Marcq et al. (2005; 2006) reported an OCS mixing ratio ~0.55  0.15
545
ppm at ~36 km, decreasing with altitude with a gradient of -0.28 ppm/km based on the
546
ground-based telescope IRTF observation. The VIRTIS measurement (Marcq et al.,
547
2008) found the OCS mixing ratio ranging between 2.5  1 and 4  1 ppm at 33 km,
548
agreeing with the previous value ~4.4 ppm from Pollack et al. (1993). Therefore, OCS
18
549
would only be ~0.1 ppm or less in the lower cloud layer (~47 km). However, a sensitivity
550
study of model A (Fig. 13) shows that 0.3-5 ppm OCS at the upper cloud deck (~58 km)
551
is required to reproduce the OCS mixing ratio at 65 km in the observed range 0.3-9 ppb
552
reported by Krasnopolsky (2010b). Krasnopolsky (2008) reported even larger values, ~14
553
ppb around 65 km and ~2 ppb around 70 km. Venus Express results suggest that the
554
upper limit of OCS is 1.6 ± 2 ppb in the region 70-90 km (Vandaele et al., 2008), and the
555
one tentative detection of OCS reported from ground-based observations near 12 m
556
(Sonnabend et al., 2005) found an abundance consistent with that calculated by Mills
557
(1998), which specified 0.1 ppm OCS at 58 km. But model A can produce only several
558
tens of ppt OCS around 70 km. Besides, the scale height of OCS in model A is ~1 km at
559
65 km, which matches only the lower limit of the observations (1-4 km from
560
Krasnopolsky, 2010b). It seems that the eddy transport in the upper cloud region might
561
not be efficient to transport OCS upward. This is also consistent with the ~4 km scale
562
height of SO2 around 68 km in the Venus Express measurements. Although the eddy
563
diffusivity at ~60 km has been estimated to be less than 4.0×104 cm2 s-1 (Woo and
564
Ishmaru, 1981), it could have large variations in the cloud layer, leading to large variation
565
in the detected OCS values and maybe related to the decadal variation of SO2 at the upper
566
cloud top (see Figure 7 of Belyaev et al., 2008).
567
568
The unexpected large amount of OCS will affect the polysulfur production. There are two
569
pathways to produce atomic sulfur (see Fig. 9). One is the photodissociation of SO and
570
ClS (Mills and Allen, 2007), The other way is from the photolysis of OCS. If the OCS
571
abundance is large (e.g. model A), the primary source of atomic sulfur below ~62 km is
572
from the photolysis of OCS instead of that of SO and ClS. There are also two pathways to
573
produce S2. One is from the chlorosulfane chemistry (Mills and Allen, 2007) and the
574
other is from the reaction between atomic sulfur and OCS. It turns out that the reaction
575
rate of S + OCS in model A is as large as the ClS2 photolysis below 60 km. Therefore if
576
there is an abundant OCS layer near the lower boundary, it may greatly enhance the
577
production of Sx in the 58-60 km region.
578
579
Elemental sulfur Supersaturation
19
580
581
Even using the highest nucleation rate ( = 1), the model A results show that the S2, S3,
582
S4 and S5 are highly supersaturated (see Fig. 6). The column abundances of gaseous S2,
583
S3, S4 and S5 above 58 km are 1.4×1013 cm-2, 9.0×1010 cm-2, and 2.1×1011 cm-2, 4.3×109
584
cm-2, respectively. S5 is supersaturated with the saturation ratio ~1000 peaking at 60 km
585
but decreases rapidly. The saturation ratio of S4 is about 107 at the lower boundary and
586
becomes moderately supersaturated above 76 km. S3 is oversaturated by a factor of 105-
587
1010 from 58 to 100 km. S2 is extremely supersaturated at all altitudes. The saturation
588
ratio is 108 at the bottom and the top, with a peak of 1015 at 90 km, where the
589
heterogeneous nucleation of S2 is negligible compared with the production processes
590
from atomic sulfur through the three-body reaction 2S + M, and the major loss processes
591
of S2 are oxidization to SO and the photo-dissociation to atomic sulfur. As illustrated in
592
Fig. 9, the main production processes of Sx can be summarized as S + Sx-1 → Sx and S2 +
593
Sx-2 → Sx, but the reactions ClS + S2 and S2O + S2O are also important for S3 production
594
at the bottom and top of the model atmosphere, respectively. The loss mechanisms of Sx
595
include the heterogeneous nucleation, conversion to other allotropes, and oxidization
596
through Sx + O → Sx-1 + SO. Fig. 15 shows the comparison of the diurnally-averaged
597
photolysis timescales of S2, S3 and S4 with those of the nucleation and eddy transport.
598
The S2 loss process is dominated by the nucleation from 58 km to about 72 km where the
599
photolysis is as fast as the nucleation, but the conversion from S2 to S4 is also important
600
around 60 km. The photolysis timescales of S3 and S4 are of the order of 1s, much smaller
601
than the nucleation timescale (~10 s at 60 km and ~100 s at 70 km). Therefore, for S 3 and
602
S4, photolysis by visible light is the major loss pathway and the heterogeneous nucleation
603
processes are negligible. Since the S3 and S4 aerosols are the possible candidates of the
604
unknown UV absorbers although they are unstable (Carlson, et al., 2010), the condensed
605
S3 and S4 are probably produced from the heterogeneous Sx chemistry over the H2SO4
606
droplet surfaces (Lyons, 2008). A proper treatment of the microphysical processes
607
coupled with atmosphere dynamical processes within the cloud layer is needed to
608
elucidate the Sx chemistry.
609
[Figure 15] (see page 57)
610
20
611
Alternative Hypotheses
612
613
Eddy diffusion is able to transport the species only from a region of high mixing ratio to
614
that of a low mixing ratio, and so it cannot generate an inversion layer. A sudden large
615
injection of SO2 from either volcano (Smrekar et al., 2010) or the instability in the cloud
616
region (VMC measurements from Markiewicz et al., 2007) may provide a sulfur source
617
at ~70 km, where the long-term natural variability of SO2 has been documented but not
618
understood. However, it is difficult for these mechanisms to explain the SO2 inversion
619
layer above 80 km because: (1) Volcano eruption could only reach 70 km but not higher
620
based on a recent Venus convective plume model (Glaze et al., 2010); (2) Even if the
621
sudden injection reaches ~100 km high, it’s also difficult to maintain the steady SO2
622
inversion for an extended period in the Venus Express era because the gas-phase SO2
623
lifetime is short (~a few earth days or less above 70 km). A continuous upwelling of SO2
624
from the lower region to the upper region (advection) is possible although the dynamics
625
maintaining the inversion profiles is not understood. The upward flux can be estimated by
626
balancing the downward flux by diffusion:
  K zz [M ]
627
df
dz
(8)
628
where  is the vertical flux, Kzz eddy diffusivity, [M] number density, f mixing ratio,
629
and z altitude. Assuming the Kzz~106 cm2s-1, df ~ 10-7, [M]~ 1017 cm-3, dz~10 km (80-90
630
km), we obtain  ~ 1010 cm-2s-1. The same magnitude of an upward flux at 80 km is
631
required to balance it.
632
633
Compared with the dynamical mechanism which transports SO2 directly from 80 km, our
634
mechanism is more plausible because: (1) the inversion can be explained by the shape of
635
the equilibrium vapor pressure profile above 90 km (models A and B); (2) it needs the
636
upward transport of aerosols only around the 90 km region, where the SSAS circulation
637
is strong and has been verified by the nighttime warm layer (Bertaux et al, 2007).
638
639
5: Summary and Conclusion
640
21
641
This study is motivated by the recent measurements from Venus Express, especially the
642
SO2 profile from 70 to 110 km and the SO profile above 80 km, and some ground-based
643
observations (SO and SO2 from Sandor et al., 2010; OCS from Krasnopolsky, 2010b).
644
The three primary chemical cycles: oxygen cycle, chlorine cycle and sulfur cycle are
645
closely coupled in the cloud region. We included the heterogeneous nucleation of
646
elemental sulfur and found that the S2, S3, S4 and S5 near the lower boundary are highly
647
supersaturated, even using the fastest removal rates by nucleation. Chlorosulfanes
648
chemistry may play an important role in producing the polysulfurs in the upper cloud
649
layer. However, in order to reproduce the recent ground-based observations, the required
650
OCS mixing ratio at the lower boundary (1.5 ppm) is found to be significantly larger than
651
the previous estimations (e.g. Yung et al., 2009). This enhanced OCS layer near 60 km
652
would greatly increase the polysulfur production rate through the photolysis to atomic
653
sulfur. But it is also possible to reduce the required OCS abundance at the lower
654
boundary if we increase the eddy diffusion transport in the upper cloud layer.
655
656
In the region above 80 km, we propose two possible solutions to explain the inversion
657
layer of SO2. The essence of our idea to solve this problem is to ‘reverse’ the sulfur cycle
658
in the region below 80 km. In 58-80 km, the SO2 and OCS are the parent species
659
transported from the lower atmosphere as the sulfur source, while the H2SO4 and possible
660
polysulfur aerosols are considered to be the ultimate sulfur sink and their production rates
661
are shown in Fig. 14. However, in the upper region the aerosols might become the sulfur
662
source rather than the sink to provide enough sulfur for the inversion layers but it requires
663
large aerosol evaporation. We relate this possible evaporation to the warm layer above 90
664
km in the night side observed by Venus Express (Bertaux et al., 2007). Therefore the SO
665
and SO2 inversion layers above 90 km are the natural results of the temperature inversion
666
induced by the adiabatic heating of the SSAS flow. While the inversion layers in the
667
region between 80-90 km are due to the downward diffusion from the lower
668
thermosphere.
669
670
If H2SO4 aerosol is the source, the cross section of H2SO4 photolysis and the SVP is
671
needed to be determined accurately. However, the laboratory work has yet to be done.
22
672
From the modeling results, the possible solutions are:
673
674
(1) Use the cross sections from Lane et al. (2008) for the UV region and Mills et al.
675
(2005) and Feierabend et al. (2006) data for the visible region, but the H2SO4
676
saturation ratio is about ~100 under nighttime temperature. That means the large
677
supersaturation exists not only in the dayside but also in the nightside. The
678
photolysis coefficient at 90 km is ~7.3×10-8 s-1.
679
680
(2) Use the UV cross sections from Lane et al. (2008) and H2SO4H2O cross sections
681
in the visible region from Vaida et al. (2003). This case requires that the hydrate
682
abundance be roughly the same order of magnitude of the pure H2SO4 saturated
683
vapor abundance. The photolysis coefficient at 90 km is ~6.8×10-6 s-1.
684
685
(3) Use the same cross sections as (1) but also use 1×10-21 cm2 molecule-1 in the UV
686
region of 195-330 nm, as shown by the dashed line in Fig. A7. The required
687
H2SO4 saturation ratio is ~0.25 under nighttime temperature. The photolysis
688
coefficient at 90 km is ~7.0×10-6 s-1.
689
690
The major difference between (3) and the other two possibilities is that, in (3) the
691
photolysis rate is contributed mainly by the UV flux, while in (1) and (2) the dominant
692
sources are the visible and IR photons.
693
694
In model A we discussed possibility (3), and the model B considers the Sx aerosol as the
695
source instead. The models A and B show some similar behaviors, which represent the
696
general features of the upper region chemistry despite the different sulfur sources.
697
Although there are uncertainties of the model parameters, the SO2 mixing ratio merely
698
depends on the input sulfur flux: in terms of the H2SO4 photolysis production rate in
699
model A and the S8 oxidization rate in model B, respectively. Because of the existence of
700
the fast sulfur cycle, we consider all the sulfur oxides in the upper region as a box, and
701
the required sulfur flux inputs in the box above 90 km is ~5×108 cm-2 s-1, consistent in
702
both models A and B. All the sulfur oxides output from models A and B, except SO3, are
23
703
very similar. This is because that the gas phase sulfur chemistry in the upper region is
704
simpler than that in the lower region (below 80 km) because it is driven by the photolysis
705
reactions and backward recombination with O and O2. However, the complexity comes
706
from the coupling of the gaseous chemistry with the aerosol microphysics and the SSAS
707
and zonal transport, both of which are poorly determined at this time. Future observations
708
and more complete modeling work are needed to fully reveal the behavior of the coupled
709
system.
710
711
Here, we present several basic differences between model A and B for future
712
consideration.
713
714
First, the two mechanisms are probably applied to different regions. By roughly
715
estimating the chemical timescales and dynamical timescales, we found that the Sx + O
716
reaction in model B is much faster than the transport. As the Sx aerosols are evaporated in
717
the nightside, the Sx vapor will be oxidized in less than 10 s, therefore the SOx is actually
718
first produced around the anti-solar region and then transported to the dayside by the
719
zonal wind and photodissociated. On the other hand, H2SO4 photolysis has to happen in
720
the dayside in the model A. So the SOx should be first produced in the dayside and then
721
transported to the nightside. A big issue of the model A is the H2SO4 condensation rate in
722
the dayside since it is highly supersaturated. If the homogenous nucleation were very fast,
723
the supersaturated vapor pressure could not be maintained, and the production of SO2
724
from the photolysis of H2SO4 would be too small. Therefore, the nightside H2SO4
725
abundances would be also supersaturated in order to supply enough H2SO4 for the
726
dayside.
727
728
Secondly, the two mechanisms might require different aerosol flux from below. Since the
729
aerosols cannot be fully recycled above 90 km due to diffusion loss of sulfur, an upward
730
aerosol flux is needed. The estimated flux is ~1 cm-2 s-1 and ~3 cm-2 s-1, corresponding to
731
an effective upward transport velocity ~0.1 cm s-1 and ~30 cm s-1 for model A and model
732
B, respectively. However, the estimation of Sx transport velocity here is based on the
24
733
assumption of the Sx/H2SO4 ratio ~1% (Carlson et al., 2010), which remains to be
734
confirmed by future measurements.
735
736
Thirdly, in terms of the possible observational evidence, the model A requires the H2SO4
737
number density ~108 cm-3 around 96 km in the dayside, which might be observed in the
738
future. But the estimated abundance of S8 in the nightside is only ~102 cm-3 around 96
739
km, which would be hard to observe. SO3 might provide another possibility to distinguish
740
the two mechanisms because SO3 is controlled mainly by the H2SO4 photolysis in model
741
D but by the SO2 oxidization in model B. The abundance of SO3 at 96 km is ~3.3×107
742
cm-3 and ~2.2×105 cm-3 for models A and B, respectively. Future observations might be
743
able to detect this difference. On the other hand, if the SO radical is produced mainly by
744
the polysulfur oxidization in the nightside, it might be possible to observe the nightglow
745
of SO excited states, in analogous to the O2 nightglow. The electronic transition of the
746
SO (a1∆→X3∑) at 1.7 μm has been detected in Io’s atmosphere (de Pater et al., 2002).
747
748
Future work is needed to advance our understanding of Venus. We briefly summarize the
749
important tasks as follows.
750
751
(1) Observations of abundances of SO3 and H2SO4 in the lower thermosphere and SO
752
nightglow in the nightside and better constraints of SO2 and OCS abundances in
753
the upper cloud region.
754
755
756
757
(2) Measurements of photodissociation cross section of H2SO4 in the UV region,
especially in the lower energy range (~195-330 nm).
(3) Laboratory measurements of H2SO4 SVP in the lower temperature region (~150240 K).
758
(4) Determination of polysulfur reaction coefficients.
759
(5) An explanation of the longtime variation of SO2 at the cloud top ~70 km and the
760
possible variation of the OCS abundance in the upper cloud region.
761
(6) The microphysical properties of Sx and H2SO4 aerosols, their formation and loss
762
processes, transport, vertical profiles and the cause of the multi-modal
763
distributions.
25
764
(7) Coupled mesosphere-thermosphere (~58-135 km) chemistry including the neutral
765
species and ions to reveal the role of SSAS transport in both dayside and
766
nightside of Venus.
767
768
Acknowledgements
769
770
We thank A. P. Ingersoll and D. Yang for insightful comments, M. Gerstell, M. Line and
771
other members of Yung’s group at Caltech for reading the manuscipt. We are indebted to
772
T. Clancy for pointing out the importance of Sx aerosols as a potential source of SO2 in
773
the mesosphere of Venus. This research was supported by NASA grant NNX07AI63G to
774
the California Institute of Technology. M. C. Liang was funded by NSC grant 98-2111-
775
M-001-014-MY3 to Academia Sinica. F. P. Mills was supported by grants under the
776
Australian Research Council Discovery Projects and Linkage International schemes. D.
777
A. Belyaev acknowledges support from CNES for a post-doc position at LATMOS.
778
779
Appendix A
780
781
Radiative Transfer
782
783
The diurnally-averaged radiation calculation here is modified on the basis of Mills
784
(1998). The direct attenuated flux and Rayleigh scattering calculations remain the same
785
(see details in the Appendix H, and I of Mills, 1998). In this study we adopt 550 log-
786
linear optical depth grids, 112 wavelengths from 960-8000 Å, 14 zenith angles for the
787
incoming photons, 8 Gaussian angles for the diffused photons. The wavelength-
788
independent middle cloud albedo at the lower boundary is assumed to be 0.6. The
789
depolarization factor of CO2 Rayleigh scattering equals 0.443.
790
791
The absorption of the unknown UV absorber and scattering processes of haze and cloud
792
particles are crucial for the radiation field, especially in the upper cloud layer. We follow
793
the procedure described in Crisp (1986). First, we calculated the optical depths from the
794
bimodal aerosol profiles (see Appendix B) and scaled to match the optical depth values in
26
795
their table 2 (equatorial cloud model) in Crisp (1986). Aerosol optical properties are
796
calculated using Mie-code based on the parameters of equator hazes in Table 1 of Crisp
797
(1986). For mode 1, the refractive index is 1.45, radius 0.490.22 m. For mode 2, the
798
refractive index is 1.44, radius 1.180.07 m. Fig. A1 shows the scattering efficiencies
799
(upper panel) and asymmetry factors (middle panel) of the two modes. Since the
800
asymmetry factors do not vary significantly, we choose 0.74 as the mean value for all
801
wavelengths. The UV absorber is introduced by decreasing the single scattering albedo of
802
the mode 1 aerosol between 3100-7800 Å. We take the empirical absorption efficiency
803
values from the Table 4 of Crisp (1986). Fig. A1 (lower panel) shows the single
804
scattering albedo of the mode 1 aerosol mixed with the UV absorber. Because the single
805
scattering albedo is not constant (from 0.85 to 1) with wavelength, we use the
806
wavelength-dependent values in the calculation. The spectral actinic fluxes (in units of
807
photons cm-2 s-1 Å-1) for 58, 62, 70, 90 and 112 km at 45N are plotted as functions of
808
wavelength in Fig. A2, although in this study our calculation is at 70N. Due to
809
absorptions by CO2, SO2 and SO, the UV flux decreases rapidly as they penetrate deeper
810
into the atmosphere. Rayleigh scattering and aerosol scattering result in the larger actinic
811
flux in the cloud and haze layers than that at the top of the atmosphere. In the wavelength
812
range large than 2000 Å, the actinic flux peaks around ~65 km. The UV actinic flux at
813
the lower boundary (~58 km) between 2000-3000 Å is roughly anti-correlated with the
814
SO2 cross sections and the minimum in its the cross section profile near 2400 Å may
815
open a window for the UV flux to penetrate to the lower atmosphere of Venus. There is
816
an analogous spectral window in the terrestrial atmosphere between 200 and 220 nm
817
(Froidevaux and Yung, 1982).
818
[Figure A1] (see page 58)
819
[Figure A2] (see page 59)
820
821
Appendix B
822
823
Nucleation Rate of Elemental Sulfur
824
825
Aerosol profile
27
826
827
Above the middle cloud top (~58 km), the aerosols are found to exhibit a bimodal
828
distribution in the upper cloud layer (58-70 km) and upper haze layer (70-90 km). In this
829
study we combine the upper haze profiles from Wilquet et al. (2010) above 72 km and
830
upper cloud particle profiles from Knollenberg and Hunten (1980) from 58-65 km. Due
831
to the lack of data for the intermediate altitudes (65-72km) at present, interpolation is
832
applied. Fig. A3 shows the bimodal aerosol profiles (left panel). From the aerosol
833
abundances, we can estimate the sulfur content. Mode 1 aerosols are roughly 0.2 μm in
834
radius constantly for all altitudes. For mode 2 aerosols, we use 0.7 μm above 72 km
835
(Wilquet et al. 2010) for the haze particles and 1.1 μm below for the cloud particles
836
(Knollenberg and Hunten, 1980). The right panel in Fig. A3 shows the equivalent sulfur
837
mixing ratio (ESMR) by volume computed from the H2SO4 aerosol (solid line)
838
abundances by assuming that the H2SO4 aerosol density is 2 g cm-3 and weight percent
839
are 85% and 75% below and above 72 km, respectively. The ESMR in the H2SO4 droplet
840
is close to 1 ppm at all altitudes, which is enough for the enhancement of sulfur oxides
841
above 80 km. By assuming that the radius of elemental sulfur is about half of the H2SO4
842
aerosol radius and the density is also 2 g cm-3, we found the ESMR in elemental sulfurs
843
in excess of 1 ppb level at most altitude (Fig. A3, right panel, dashed line).
844
[Figure A3] (see page 60)
845
846
Heterogeneous Nucleation
847
848
The nucleation rate of elemental sulfurs onto H2SO4 droplets is estimated as follows. The
849
nucleation rate constant in the continuum regime (where the particle size is much larger
850
than the vapor mean free path ) is expressed as (Seinfeld and Pandis, 2006):
851
J c  4 Rp Ds , where Rp is the H2SO4 aerosol radius and Ds the molecular diffusivity of
852
elemental sulfur vapor.
853
854
However, in the Venus cloud layer, the Knudsen Number Kn = /Rp of Sx vapor is not far
855
from 1, so the nucleation process lies in the transition regime where the mean free path 
856
of the diffusing vapor molecule (e.g., Sx vapor) is comparable to the pre-existing aerosol
28
857
size. Therefore, we adopt the Dahneke approach (Dehneke, 1983) modified to include the
858
molecular accommodation coefficient (Seinfeld and Pandis, 2006). The approach
859
matches the fluxes of continuum regime ( K n = 1 ) and free molecular regime ( K n ? 1 )
860
by introducing a function f (K n ) :
f (K n ) 
861
1  Kn
1  2K n (1  K n ) / 
(A1)
862
where  is the molecular accommodation coefficient, which is the probability of sticking
863
when the vapor molecule encounters a particle. Here the mean free path  in Kn is
864
defined as 2Ds/v, where v is the mean thermal velocity of the vapor molecule.
865
866
Finally we obtain the nucleation rate constant:
J  f (K n )J c 
867
4 Rp Ds (1  K n )
(A2)
1  2K n (1  K n ) / 
868
The molecular diffusivity Ds of sulfur vapor can be estimated using hard sphere
869
approximation: Ds=b/N, where N is the total CO2 gas density in the environment and b is
870
the binary collision parameter:
 2 kT (ms  mg ) 
3
b

4 (ds  dg )2 
ms mg

871
1/2
(A3)
872
where ds and dg are the diameters of sulfur vapor and CO2 gas molecule, respectively
873
(assuming ds = dg= 3 Å), k = Boltzmann constant, T = temperature, ms and mg are the
874
mass of sulfur vapor and CO2 gas molecule, respectively. Fig 15 shows the total
875
nucleation timescale (from two modes of aerosols) of S2 (roughly the same as other
876
allotropes), together with the eddy transport timescale and photolysis timescales of S2, S3
877
and S4. See the discussion in section 4.
878
879
Appendix C
880
881
H2SO4 and Sx vapor abundances
882
883
H2SO4
29
884
885
If sulfuric acid is in thermodynamic equilibrium with the surrounding atmosphere, the
886
saturation vapor pressure (SVP) over H2SO4 aerosol should mainly depend on the
887
temperature and aerosol composition. However, non-thermodynamic equilibrium in the
888
real atmosphere is common because the chemical and dynamic processes, such as the
889
chemical production, condensation and transport, are often involved and play important
890
roles. The condensation efficiency, which depends on many microphysical properties of
891
the system like the temperature, diffusivity, aerosol size, surface tension, and interaction
892
between molecules and aerosols, will greatly affect the H2SO4 vapor pressure over the
893
liquid droplets. The very low condensation rate could cause large supersaturation of the
894
H2SO4 vapor. For example, the saturation ratio of H2SO4 in the lower stratospheric sulfate
895
layer (Junge layer) on Earth has been observed to be as large as 102-103 (Arnold, 2006).
896
A similar situation may exist in the Venus upper haze layer on the dayside when the
897
H2SO4 vapor on the night side is transported to the dayside, because the SVP of H2SO4 in
898
the night side is several orders of magnitude larger than that in the dayside (Zhang et al.,
899
2010) due to the large temperature difference above 90 km. Zhang et al. (2010) proposed
900
that this might be the key mechanism to explain the SO2 inversion layer because the
901
nighttime H2SO4 abundance could be enough to produce the observed SO2 under
902
photochemical processes if the H2SO4 photolysis cross section is 100 time larger than the
903
current data from Vaida et al. (2003).
904
905
In the condensation processes, we assume that the sulfate aerosol will quickly establish
906
equilibrium with respect to water because there are more collisions of aerosol particles
907
with H2O molecules than with H2SO4 molecules. Therefore, we could derive the H2SO4
908
aerosol composition (weight percent) from the water activity (or equilibrium relative
909
humidity) defined as the partial pressure of water vapor divided by the SVP over pure
910
water under the same temperature. The water activity is shown in Fig. A4 (left panel) for
911
day and night temperature profile, respectively. We used the H2O SVP as function of
912
temperature from Tabazadeh et al. (1997), which is valid between 185-260K.
913
PH 2 O
3.5052  10 3 3.3092  10 5 1.2725  10 7
 exp(18.4524 


)
T
T2
T3
30
(A4)
914
where PH 2 O is the SVP of H2O in mbar and T is temperature. We extrapolated the formula
915
to the entire temperature range (156-274 K) of Venus mesosphere so there would be
916
some uncertainties above 84 km for the dayside temperature and in the 84-90 km for the
917
nighttime temperature.
918
[Figure A4] (see page 61)
919
920
The H2SO4 weight percent is therefore roughly estimated by comparing the observed H2O
921
mixing ratio profile with the theoretical profiles under different H2SO4 compositions, as
922
shown in Fig. A4 (the middle and right panels). For the 50-80 wt% H2SO4, we used the
923
table from Tabazadeh (1997), calculated based on the Clegg and Brimblecombe (1995).
924
For the more concentrated acids, our calculation is based on Gmitro and Vermeulen
925
(1964) although it may not be very accurate for the low temperature (Mills, 1998). There
926
are also some uncertainties in applying the Tabazadeh (1997) formula to Venus because
927
the Clegg and Brimblecombe (1995) is only valid if the water activity larger than 0.01.
928
The atmosphere of Venus is very dry (Fig. A4), and so actually only the results in the
929
region from 85-100 km in the dayside and 85-90 km in the nightside seem robust.
930
However, as we show in the last paragraph, the H2O SVP may have some uncertainties in
931
those regions. Therefore, the H2SO4 weight percent derived here is only a rough estimate
932
based on the current knowledge.
933
934
The H2SO4 weight percent falls with altitude, associated with the increase of relative
935
humidity due to the temperature decrease. The values are about 90%-84% in 58-70 km
936
and 84%-60% in 70-90 km, which are roughly consistent with the H2SO4 compositions
937
obtained from aerosol refractive indexes based on the photometry measurements (85%
938
and 75%, respectively). But in the region above 90 km, the large contrast of dayside and
939
nightside temperatures results in large difference of the local H2SO4 weight percent. For
940
example, H2SO4 at 100 km is ~75% in the dayside but can be larger than 96% in the
941
nightside. Actually the temperature profile above 90 km has been found to be a function
942
of longitude (Bertaux et al., 2007). Therefore, if the transport is efficient, the H2SO4
943
aerosols could have a broad range distribution of various concentrations above 90 km but
31
944
the H2SO4 vapor abundances might be mainly determined by the warmest nightside
945
temperature since the vapor abundances is extremely sensitive to the temperature.
946
947
The H2SO4 SVP is another uncertainty and maybe the major one. In the supplementary
948
material of Zhang et al. (2010), three H2SO4 SVP formulas as function of temperature
949
and H2SO4 concentration have been discussed in details. These formulas could differ by
950
several orders of magnitude but none of them has been verified in the temperature range
951
of upper atmosphere of Venus. Instead of using the H2SO4 weight percent profile derived
952
in Fig. A4, we simply assumed 85% H2SO4 below 70 km and 75% from 70-90 km and
953
used the vapor pressure formulas from Ayers et al. (1980) corrected by Kulmala and
954
Laaksonen (1990):
955
ln P 0H 2 SO4  16.259 
 1
0.38 
T0 T0  
 10156    
 1  ln( )   
8.3143T
T
T 
 T Tc  T0
  0
(A5)
956
where Tc = 905 K, T0 = 360.15 K, P0H2 SO4 is SVP of H2SO4 in atm, T is the temperature, 
957
and 0 are the chemical potentials of H2SO4 solutions of certain composition and pure
958
acid, respectively. The values of -0 for the 85% and 75% H2SO4 are 1555 cal-1 mole
959
and 3681 cal-1 mole based on Giauque et al. (1960), respectively.
960
961
In fact the H2SO4 abundances in the region below 80 km are not important because the
962
H2SO4 photolysis is negligible for the lower region chemistry. But in the upper region the
963
H2SO4 might behave like a sulfur source rather than a sink, and large abundance of
964
H2SO4 is required in the upper region in order to reproduce the SO2 inversion layer
965
(Zhang et al., 2010). So we adopted the formula by Stull (1947) just for reference, simply
966
because it gives the largest SVP in the Venus temperature range:
967
PH2 SO4  103954.90/T 9.4570
968
where PH 2 SO4 is the SVP of H2SO4 in mmHg and T is temperature. The H2SO4 SVP
969
profiles in Fig. A5 (left panel) show large difference between the dayside and nightside
970
caused by the difference in temperatures. Since H2SO4 is very hygroscopic, the right
971
panel shows the abundance of monohydrate (H2SO4H2O), estimated based on the
972
extrapolation of the equilibrium constants from the Vaida et al. (2003) for the earth
32
(A6)
973
atmosphere (223-271 K in the literature). The abundances of H2SO4H2O above 90 km
974
are less than 5% and much less (<10-5) of that of pure H2SO4 for the dayside and
975
nightside, respectively, although the equilibrium constants have not been verified in the
976
Venus temperature region (~160-240 K).
977
[Figure A5] (see page 62)
978
979
Sx
980
981
Lyons (2008) summarized the previous laboratory measurements and computed vapor
982
pressure over the liquid sulfur allotropes and the total solid sulfur vapor pressures over
983
the orthorhombic sulfur and the monoclinic sulfur below the melting points. Based on the
984
data, he estimated the equilibrium vapor pressure over solid sulfur allotropes. In this
985
study, we follow the same method and calculated the monoclinic Sx saturated volume
986
mixing ratio profiles under the daytime and nighttime temperature situations. The results
987
are shown in Fig. A6.
988
[Figure A6] (see page 63)
989
990
Appendix D
991
992
H2SO4 photolysis cross section
993
994
H2SO4 was thought to be photodissociated by the UV photons only. Burkholder et al.
995
(2000) and Hintze et al. (2003) estimated the upper limits for the UV cross section of
996
H2SO4 based on the failure to detect any absorption beyond 140 nm. The upper limits are
997
assumed to be 1×10-21 cm2 molecule-1 in the interval 195-330 nm, 1×10-19 cm2 molecule-1
998
in 195-160 nm, and 1×10-18 cm2 molecule-1 in 160-140 nm. Lane et al. (2008) revisited the
999
UV cross sections by calculating the electronic transitions based on the theoretical twin
1000
hierarchial approach and they found that the cross section in the Lyman- region (~121.6
1001
nm) is about ~6×10-17 cm2 molecule-1, much larger than the previously assumed value.
1002
And it also seems that the cross section in the interval 195-330 nm is much smaller than
1003
the upper limits from Burkholder et al. (2000).
33
1004
1005
Vaida et al. (2003) proposed that in the visible region the excitation of the OH-stretching
1006
overtone transitions with   4 (~38.6 kcal mole-1, or ~742 nm) is also enough to
1007
photolyze H2SO4 because the energy required for H2SO4 + hυ → SO3 + H2O is only 32-40
1008
kcal mole-1. This mechanism has been verified by the laboratory experiments in 49 and
1009
59 bands from the cavity ring-down spectroscopy by Feierabend et al. (2006). Vaida et al.
1010
(2003) also proposed that, in the IR and visible regions the OH-stretching overtone
1011
transitions with   3 (~26.3 kcal/mole, or ~1.09 m) are able to generate the
1012
photodissociation of H2SO4H2O as well (required energy ~25 kcal mole-1) and the total
1013
photolysis coefficient is about ~100 times larger than that of pure H2SO4, although a recent
1014
simulation by Miller et al. (2006) suggested that the H2SO4H2O is more likely to
1015
thermally decompose to H2SO4 and H2O before photodissociation.
1016
1017
The solid line in Fig. A7 shows the cross sections from Lane et al. (2008) for the UV
1018
region and Mills et al. (2005) and Feierabend et al. (2006) data for the visible region and
1019
binned in our model spectral grid. As shown in Table 1, the H2SO4 photolysis coefficient is
1020
generally ~10-7 s-1 in the upper atmosphere. It is ~10-6 s-1 near the upper boundary (112
1021
km) due to the photolysis by the Lyman- line but only in a very thin layer (<1 km)
1022
because the Layman alpha intensity decreases very rapidly due to absorption by the CO2.
1023
The major contribution to the photolysis is the solar pumping of the vibrational overtones
1024
by the 740 nm red light (49 band, Vaida et al., 2003). The collisional deactivation mainly
1025
depends on the atmospheric pressure. In Miller et al. (2007) the quantum yield is nearly
1026
unity above 60 km where the pressure is 0.2 mbar in the Earth atmosphere. In Venus, this
1027
pressure level (0.2 mbar) is at ~90 km which is the lower boundary of the region we are
1028
interested here. Therefore the quantum yield is assumed to be unity above 90 km.
1029
[Figure A7] (see page 64)
1030
1031
However, Zhang et al. (2010) showed that the photolysis coefficient ~10-7 s-1 is not
1032
enough to produce the observed SO2, otherwise a very large supersaturation of H2SO4
1033
(~100) under nighttime temperature is needed. Although this supersaturation is possible
1034
(as seen in Earth), empirically they also found the required cross section is about ~100
34
1035
times larger than that of pure H2SO4 if keeping the H2SO4 vapor abundances roughly the
1036
same as the nighttime saturated abundances. This extreme situation may suggest the
1037
existence of large amount of H2SO4H2O and maybe other hydrates (like H2SO42H2O),
1038
although it seems not very likely not only because the equilibrium abundance of the
1039
monohydrate is small (see Fig. A5) but also because the sulfuric acid hydrates might
1040
readily condense into the crystal phase even under the nighttime temperature
1041
(McGouldrick et al., 2010). Alternatively, the required large cross section actually could
1042
be achieved by assuming the UV cross section as the upper limit of 1×10 -21 cm2
1043
molecule-1 between 195 and 330 nm, as shown in dashed line in Fig. A7. We consider this
1044
possibility in the model A. Note that this change of H2SO4 photolysis may not affect
1045
much for the earth stratosphere below 35 km because of the absorption of O3 Hartley
1046
band dominates the actinic flux in that region. However, this is very important for the
1047
Venus mesosphere above the cloud top since the SO2 absorption is not as strong as O3.
1048
The H2SO4 photolysis coefficient in this case is ~8.3×10-6 s-1 at 90 km, roughly the same
1049
as that of ~8.2×10-6 s-1 if we use the H2SO4H2O photolysis cross section instead (model
1050
B in Zhang et al., 2010).
1051
1052
References
1053
1054
1055
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possible influence of cosmic rays?, Space Sci. Rev. 125, 169-186.
1056
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1057
chemistry: Volume I-gas phase reactions of Ox, HOx, NOx and SOx species,
1058
Atmos. chem.phys. 4, 1461-1738.
1059
1060
[3]Ayers, G. P., et al., 1980, On the vapor pressure of sulfuric acid, Geophys. Res. Lett.
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1061
[4]Bahou, M., 2001, Absorption cross sections of HCl and DCl at 135-232 nanometers:
1062
Implications for photodissociation on Venus, Astrophys. J. Lett. 559, L179.
1063
[5]Baulch, D. L., et al., 1992, Evaluated Kinetic Data for Combustion Modelling, J. Phys.
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Chem. Ref. Data 21, 411-734.
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1065
[6]Baulch, D. L., et al., 1981, Evaluated Kinetic Data for High Temperature Reactions
1066
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1067
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measurements of HF, HCl, H2O and HDO, Nature 450, 646-649.
1074
[10]Brune, W. H., 1983, Laser magnetic resonance, resonance fluorescence, and
1075
resonance absorption studies of the reaction kinetics of O+OH→H+O2, O+
1076
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1077
and 5 torr, J. phys. chem. A 87, 4503.
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[11]Bryukov, M. G., 1993, Kinetics of HNO reactions with O2 and HNO, Chem. Phys.
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1256
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42
1264
1265
Fig. 1. Left panel: daytime temperature profile (solid) and nighttime temperature profile
1266
(dashed). Right Panel: eddy diffusivity profile (solid) and total number density profile
1267
(dashed).
1268
1269
1270
1271
1272
1273
1274
1275
1276
1277
1278
1279
1280
43
1281
1282
Fig. 2. Volume mixing ratio profiles of the oxygen species from model A.
1283
1284
1285
1286
1287
1288
1289
1290
1291
1292
1293
1294
1295
1296
1297
1298
44
1299
1300
Fig. 3. Same as Fig 2, for hydrogen species.
1301
1302
1303
1304
1305
1306
1307
1308
1309
1310
1311
1312
1313
1314
1315
1316
45
1317
1318
Fig. 4. Same as Fig 2, for chlorine species.
1319
1320
1321
1322
1323
1324
1325
1326
1327
1328
1329
1330
46
1331
1332
Fig. 5. Same as Fig 2, for chlorine-sulfur species.
1333
1334
1335
1336
1337
47
1338
1339
Fig. 6. Volume mixing ratio profiles (solid) of the elemental sulfur species from model A.
1340
The dashed lines are the monoclinic Sx saturated mixing ratio profiles over solid Sx based
1341
on Lyons (2008), in equilibrium with the daytime temperatures.
1342
1343
1344
1345
1346
1347
1348
1349
1350
1351
1352
1353
1354
1355
1356
48
1357
1358
Fig. 7. Same as Fig 2, for nitrogen species. The NO measurement (5.5±1.5 ppb) is from
1359
Krasnopolsky (2006).
1360
1361
1362
1363
1364
1365
1366
1367
1368
49
1369
1370
Fig. 8. Same as Fig 2, for the sulfur oxides. The SO2 and SO observations with errorbars
1371
are from the Belyaev et al. (2010). The temperature at 100 km is 165-170 K for the
1372
observations. The OCS measurement (0.3-9 ppb with the mean value of 3 ppb) is from
1373
Krasnopolsky (2010).
1374
1375
1376
1377
1378
1379
1380
50
1381
1382
Fig. 9. Important chemical pathways for sulfur species.
1383
1384
1385
1386
1387
1388
51
1389
1390
Fig. 10. Important production and loss rates for SO (upper), SO2 (middle) and SO3
1391
(lower) from model A
1392
52
1393
1394
Fig. 11. Comparison of volume mixing ratio profiles of SO2 (left), SO (middle) and SO3
1395
(right) from models A (black) and B (red).
53
1396
1397
Fig. 12. Important production and loss rates of SO (upper), SO2 (middle) and SO3 (lower)
1398
from model B.
1399
54
1400
1401
1402
Fig. 13. Sensitivity studies based on model A. Left panel: cases for different
1403
accommodation coefficients (see labels) of the heterogeneous nucleation rate. Middle
1404
panel: cases with different boundary conditions of SO2 (see labels). Right panel: cases
1405
with different boundary conditions of OCS (see labels).
1406
1407
1408
1409
1410
1411
1412
1413
1414
55
1415
1416
Fig. 14. Production rate profiles of H2SO4 (solid) and Sx (dashed) for models A (black)
1417
and B (red).
1418
1419
1420
1421
56
1422
1423
Fig. 15. Timescales for the heterogeneous nucleation, eddy transport and photolysis for
1424
S2, S3 and S4.
1425
1426
1427
1428
1429
1430
57
1431
1432
Fig. A1. Optical Properties of mode 1 (solid) and mode 2 (dashed) aerosols above the
1433
middle cloud top (~58 km) based on the parameters from Crisp (1985). Upper panel:
1434
extinction efficiency; Middle panel: Asymmetry Factor; Lower panel: single scattering
1435
albedo mixed with the empirical albedo of the unknown UV absorber from Crisp (1985).
1436
1437
1438
1439
1440
1441
58
1442
1443
Fig. A2. Spectral actinic flux in the middle atmosphere of Venus at 45 N.
1444
1445
1446
1447
1448
1449
1450
1451
1452
59
1453
1454
Fig. A3. Left panel: Concentration profiles of mode 1 (solid) and mode 2 (dashed)
1455
aerosols based on the upper haze profiles from Wilquet et al. (2010) above 72 km and
1456
upper cloud particle profiles from Knollenberg and Hunten (1980) from 58-65 km. Data
1457
are not available in the red region (~66-70 km). Right panel: The equivalent sulfur
1458
mixing ratio (ESMR) by volume computed from the H2SO4 aerosol (solid line) and the
1459
polysulfur aerosol (dashed line). See the text for details.
1460
1461
1462
1463
1464
1465
1466
1467
1468
1469
1470
60
1471
1472
Fig. A4. Left panel: water activity profiles based on the daytime and nighttime
1473
temperature profiles. Middle and Right panels: equilibrium H2O mixing ratio contours
1474
from which H2SO4 weight percent for each altitude could be inferred by comparing the
1475
observed H2O profile (dashed line). The middle and right panels are for the day and night
1476
temperature situations, respectively.
1477
1478
1479
1480
1481
1482
61
1483
1484
Fig. A5. Left panel: H2SO4 vapor mixing ratio profiles in equilibrium with the daytime
1485
(solid) and nighttime temperatures (dashed). Right panel: monohydrate (H2SO4H2O)
1486
vapor mixing ratio profiles.
1487
1488
1489
1490
1491
1492
62
1493
1494
Fig. A6. Saturated vapor volume mixing ratio profiles of Sx allotropes based on the
1495
monoclinic Sx saturated vapor pressure over solid Sx based on Lyons (2008), in
1496
equilibrium with the daytime (dashed) and nighttime (solid) temperatures.
1497
1498
1499
1500
1501
1502
63
1503
1504
Fig. A7. H2SO4 cross sections binned in the model grids. Solid: data from Lane et al.
1505
(2008) for the UV region, and Mills et al. (2005) and Feierabend et al. (2006) for the
1506
visible region. Dashed: same as the solid line but also with 1×10-21 cm2 molecule-1 in the
1507
UV region of 195-330 nm.
1508
64