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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, E00B19, doi:10.1029/2008JE003202, 2008
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Introduction to the special section on Venus Express: Results of the Nominal Mission D. V. Titov,1,2 F. W. Taylor,3 and H. Svedhem4 Received 21 May 2008; accepted 7 August 2008; published 17 December 2008.
Citation: Titov, D. V., F. W. Taylor, and H. Svedhem (2008), Introduction to the special section on Venus Express: Results of the Nominal Mission, J. Geophys. Res., 113, E00B19, doi:10.1029/2008JE003202.
[1] More than 25 spacecraft from the United States and the Soviet Union visited Venus in the 20th century, but in spite of the many successful measurements they made, a great number of fundamental problems in the physics of the planet remained unsolved [Taylor, 2006; Titov et al., 2006]. In particular, a systematic and long-term survey of the atmosphere was missing, and most aspects of atmospheric behavior remained puzzling. After the Magellan radar mapping mission ended in 1994, there followed a hiatus of more than a decade in Venus research, until the European Space Agency took up the challenge and sent its own spacecraft to our planetary neighbor. The goal of this mission, Venus Express, is to carry out a global, long-term remote and in situ investigation of the atmosphere, the plasma environment, and some aspects of the surface of Venus from orbit [Titov et al., 2001; Svedhem et al., 2007]. [2] Venus Express continues and extends the investigations of earlier missions by providing detailed monitoring of processes and phenomena in the atmosphere and near-space environment of Venus. Radio, solar, and stellar occultation, together with thermal emission spectroscopy, sound the atmospheric structure in the altitude range from 150 to 40 km with vertical resolution of few hundred meters, revealing strong temperature variations driven by radiation and dynamical processes. In particular, by stellar occultation, Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV) discovered a warm layer at the mesopause (100 km) apparently caused by adiabatic heating of descending air. Observations by the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) and radio occultation sounding by Venus Radio Science Experiment (VeRa) indicate remarkable latitudinal temperature variations at the cloud tops, as well as the presence of a convective region within the cloud deck all over the planet. The temperature field in the mesosphere of the southern hemisphere is similar to that observed in the northern hemisphere by earlier missions, thus indicating global north/south symmetry. 1 Max-Planck Institute for Solar System Research, Katlenburg-Lindau, Germany. 2 On leave from the Space Research Institute, Moscow, Russia. 3 Department of Atmospheric, Oceanic, and Planetary Physics, Oxford University, UK. 4 ESA/ESTEC, Noordwijk, Netherlands.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003202$09.00
[3] Multispectral imaging is performed to investigate the cloud morphology at different levels and its latitudinal and temporal changes. Venus Monitoring Camera shows evidence of convection in the cloud layer at low latitudes. Streaky cloud features indicate that the convection ceases in the middle and high latitudes. Sequences of images are used to track motions of the cloud features and to derive wind speeds in the 50– 70 km altitude range. The polar orbit of Venus Express provides near-nadir viewing of middle and high latitudes in the southern hemisphere, enabling the first detailed observations of the cloud morphology and dynamics, leading to the discovery and characterization of the complex ‘‘eye’’ of the southern polar vortex by VIRTIS. The imaging instruments also perform thermal mapping of the surface in the near-IR spectral windows on the nightside, searching for surface emissivity anomalies and signs of active volcanism. [4] The chemistry and dynamics of the mesopause region (100 km altitude) is being studied by observing nonlocal thermodynamic equilibrium (non-LTE) emission from O2 and NO molecules in the UV and near-IR. This airglow peaks at the equator, close to midnight and approximately at the mesopause level, indicating a thermospheric solarantisolar component in the atmospheric circulation. Composition measurements over a wide range of altitudes are providing vertical profiles of CO, H2O, HDO, HCl, HF, and SO2 in the mesosphere (70 – 100 km), and global mapping of CO, COS, H2O, and SO2 in the lower atmosphere at heights around 35 km. These results provide powerful tests of dynamical and chemical models of the Venusian atmosphere. [5] The magnetometer (MAG) and Analyzer of Space Plasma and Energetic Atoms (ASPERA) measure the magnetic field and densities of neutral atoms, ions, and electrons in situ. These observations determine the structure and properties of the circumplanetary plasma and characterize escape processes at Venus. They cover the time of solar minimum and thus complement Pioneer Venus investigations at solar maximum. Interestingly, the H/O ratio was found to be approximately two thus suggesting that these ions are escaping in the stoichiometric ratio of water. The magnetometer is detecting whistler signals during 70% of the pericenter passes. This is interpreted as evidence of lightning and the rate is estimated to be at least as frequent as on Earth. [6] The first results from the Venus Express mission were published in nine papers in the special section ‘‘Venus
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Express’’ in Nature, 450, 629 – 662, 2007. This special section contains original papers describing results and analysis of the observations during the whole nominal mission (4 June 2006 to 2 October 2007), along with supporting modeling efforts. The papers are grouped in topical blocks and are published in two groups. This first group of papers contains a mission overview and papers on atmospheric composition, atmospheric dynamics, non-LTE emissions, and surface investigations. A second group of papers, to appear about 3 months later, will contain papers on atmospheric structure, clouds and hazes, and plasma environment.
References Svedhem, H., et al. (2007), Venus Express: The first European mission to Ve n u s , P l a n e t . S p a c e S c i . , 5 5 , 1 6 3 6 – 1 6 5 2 , d o i : 1 0 . 1 0 1 6 / j.pss.2007.01.013.
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Taylor, F. W. (2006), Venus before Venus Express, Planet. Space Sci., 54, 1249 – 1262, doi:10.1016/j.pss.2006.04.031. Titov, D. V., E. Lellouch, F. W. Taylor, and L. Marinangeli, H. Opgenoorth, and the Venus Express Team (2001), Venus Express: An orbiter for the study of the atmosphere, the plasma environment, and the surface of Venus, Mission Definition Rep. ESA-SCI (2001)6, Eur. Space Agency, Paris. Titov, D. V., H. Svedhem, and F. W. Taylor (2006), The atmosphere of Venus: Current knowledge and future investigations, in Solar System Update, edited by P. Blondel and J. W. Mason, pp. 87 – 110, Springer, Berlin. H. Svedhem, ESA/ESTEC, Keplerlaan 1, Postbus 299, NL-2200 AG Noordwijk, Netherlands. F. W. Taylor, Department of Atmospheric, Oceanic, and Planetary Physics, Oxford University, Parks Road, Oxford OX1 3PU, UK. D. V. Titov, Max-Planck Institute for Solar System Research, Max Planck Strasse 2, D-37191 Katlenburg-Lindau, Germany. (
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114, E00B33, doi:10.1029/2008JE003290, 2009
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Venus Express mission H. Svedhem,1 D. Titov,2 F. Taylor,3 and O. Witasse1 Received 30 October 2008; revised 18 December 2008; accepted 5 January 2009; published 19 March 2009.
[1] Venus Express is well and healthy and has now been providing exciting new data from
Venus, our nearby twin planet, for over 2 years. Many of the new results are presented and discussed in the subsequent papers in this special section. The overall scientific objective of Venus Express is to carry out a detailed study of the atmosphere of Venus, including the interaction of the upper atmosphere with the solar wind and the interaction of the lowest part of the atmosphere with the surface of the planet. In addition, the plasma environment and magnetic fields as well as some aspects of the surface of the planet are addressed. For the first time, investigations make systematic use of the transparent infrared spectral windows in order to probe the atmosphere in four dimensions: three spatial dimensions plus time. The spacecraft design is taken from Mars Express with some modifications necessary owing to the specific environment around Venus. The payload is composed of three spectrometers, a camera, a magnetometer, an instrument for detecting energetic particles, and a radio science package. The orbit is polar and highly elliptic, with a pericenter altitude of about 200 km over the northern polar region and an apocenter altitude of 66,000 km. Presently, the coverage of the southern hemisphere is very good, but important gaps still do exist. The coverage of the northern hemisphere is much less dense. Venus Express is a part of the European Space Agency’s program for the exploration of the inner solar system, which includes missions to study the Sun, Mercury, Venus, the Moon, Mars, and comets and asteroids. Citation: Svedhem, H., D. Titov, F. Taylor, and O. Witasse (2009), Venus Express mission, J. Geophys. Res., 114, E00B33, doi:10.1029/2008JE003290.
1. Introduction [2] Venus was a forgotten planet for more than a decade after the emphasis for investigations of the terrestrial planets shifted from Venus toward Mars during the late 1980s. There were, however, still a large number of fundamental questions to be answered about the past, present and future of Earth’s sister planet, and for an improved understanding of the general evolution of the terrestrial planets better knowledge of Venus is essential [Taylor, 2006; Titov et al., 2006a; Moroz, 2002]. The European Space Agency (ESA) launched Venus Express to open up opportunities for new investigations with a combination of instruments employing completely new techniques and improved versions of conventional instruments. The response to the initial results and findings has been very enthusiastic, generating renewed interest both in the scientific community and in the major space agencies, with new missions being developed or planned in Japan, the United States, and Russia. Venus has come back into the forefront in planetary science.
1
ESA, ESTEC, Noordwijk, Netherlands. Max-Planck Institute for Solar System Research, Katlenburg-Lindau, Germany. 3 Department of Atmospheric, Oceanic, and Planetary Physics, Oxford University, Oxford, UK. 2
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003290$09.00
[3] Venus Express is an independent follow-up to the successful Mars Express mission, launched a few years earlier, and is reusing much of the basic design for spacecraft, launcher and ground system in order to keep costs down [Titov et al., 2001; Svedhem et al., 2007a, 2009]. It is the third spacecraft in the family that started with the (much larger) Rosetta spacecraft. Both Venus Express and Mars Express benefit greatly from the generic developments carried out for that mission. With only 3 years between mission approval and the launch date, plus about a year of preparatory work preapproval, Venus Express is by far the most rapidly developed scientific project of ESA. The short developments time and the significant design heritage have enabled a powerful mission to materialize at a substantially lower cost compared to a single mission developed in isolation. [4] With the new data from Venus Express, there is a picture emerging of Venus where the existing conditions and the ongoing processes are becoming clearer. Comparisons with the Earth will continue and intensify. Venus and Earth are very different twin planets, and while they will remain different, we are moving closer toward understanding how and why the two planets have their distinct characteristics, particularly with regard to climate [Svedhem et al., 2007b]. [5] This paper describes the mission objectives of the Venus Express mission and gives an overview of the main features of the spacecraft, with emphasis on the differences compared to Mars Express, and the scientific instruments on board. It discusses the mission scenario and the operational
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orbit and the specific questions related to the choice of orbit and it gives an introduction to the various activities of the scientific operations. It discusses the coverage achieved to date by the different instruments with respect to longitude, latitude and local solar time, for the atmospheric and surface observations as well as the coverage of the different regions in and around the induced magnetosphere. Temporal coverage is discussed for a few special cases but the principles are applicable to most of the measured parameters. It outlines the place of Venus Express in the scientific program of the European Space Agency and finally it discusses topics to be studied in an extended mission, beyond May 2009, that was requested during the latter part of 2008.
2. Mission Objectives and Investigations [6] The overall scientific objective of Venus Express is to carry out a detailed and systematic study of a large number of aspects of the atmosphere of Venus. This includes the interaction of the upper atmosphere with the solar wind and the interaction of the lowest part of the atmosphere with the surface of the planet. In addition, dedicated instruments are focusing on the study of the plasma environment and magnetic fields. Some aspects of the surface are addressed. The objectives have been structured and organized under the following seven themes: (1) atmospheric structure, (2) atmospheric dynamics, (3) atmospheric composition and chemistry, (4) cloud layer and hazes, (5) energy balance and greenhouse effect, (6) plasma environment and escape processes, and (7) surface properties and geology. [7] The evolution of the planet and in particular the atmosphere and the climate are topics of great interest that span over all the themes. Most themes include aspects of comparative planetology and comparison with the other terrestrial planets, the Earth in particular, is an important objective. Direct comparison of data from Mars Express and Venus Express are enabled by the fact that several instruments are identical or similar on the two spacecraft. Also, several of the Venus Express scientists are involved in both missions. This opens up possibilities not found before on interplanetary spacecraft. 2.1. Atmospheric Structure [8] Knowledge about the atmospheric structure is of great importance for the understanding of the state of the atmosphere and the processes active therein. Previous missions provided a basic understanding of the thermal structure, but many questions remained to be answered [Taylor et al., 1980; Seiff et al., 1985]. The atmosphere can be roughly divided into three main layers based on the temperature distribution and the various processes that govern the state of the layers: the troposphere (0 –60 km), the mesosphere (60 – 100 km), and the thermosphere (above 100 km). The isothermal region between the troposphere (where temperature falls with height) and the mesosphere (where temperature increases with height) that would correspond to the stratosphere on the Earth is of very limited vertical extent on Venus and is not usually described as a separate layer. [9] Below 30 km the temperature is believed to be fairly constant all over the planet [Seiff et al., 1985] but the latitudinal coverage of the data is fairly poor. The mesosphere shows a much more variable temperature, especially in
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latitude, but what the dominating processes are is not clear [Lellouch et al., 1997]. [10] The thermosphere turned out to be surprisingly cool, especially at night (leading to the coining of the term ‘‘cryosphere’’ for the nighttime thermosphere). While it is clear that radiative processes play a major role, additional, and more densely sampled data covering all latitudes and local times are needed to achieve a good understanding [Fox and Bougher, 1991]. [11] Venus Express studies the thermal structure of the atmosphere by means of global, simultaneous, and spatially resolved spectroscopic observations in the wavelength range from UV to thermal IR. Specifically, it (1) investigates the upper atmosphere (140– 90 km) at high vertical resolution through solar and stellar occultation measurements [Bertaux et al., 2007b], (2) sounds the temperature of the middle atmosphere (100– 60 km) using spectroscopic observations of the 4.3 mm CO2 bands, with a coverage in latitude and local solar time corresponding to a spatial resolution of a few tens of kilometers or better [Formisano et al., 2006; Drossart et al., 2007a], (3) sounds the temperature in the altitude range 80– 40 km by radio occultation, providing vertical resolution of a few hundred meters [Ha¨usler et al., 2006; Pa¨tzold et al., 2007], and (4) maps the surface temperature on the nightside as a function of surface elevation [Drossart et al., 2007a]. [12] Together, these different techniques cover the range from 140 km down to 35 km. The optical measurements can be made only during nighttime since it would be very difficult to discriminate between thermal radiation and scattered sunlight. Unfortunately the PFS instrument, which would have used the 15 mm CO2 band for daytime profiling, is not operating. The use of other instruments has been rescheduled in order to as much as possible minimize the impact of this loss. 2.2. Atmospheric Dynamics [13] Data from previous missions and ground-based observations have shown that the general circulation of the atmosphere can be divided into two regimes: A retrograde zonal superrotation in the troposphere and mesosphere [Gierasch et al., 1997] and a solar to antisolar component across the terminator in the thermosphere [Bougher et al., 1997]. The zonal superrotation has a maximum wind velocity of about 100 m/s at the cloud top level (70 km), decreasing to almost zero at the surface. At the same time, there is a ‘‘Hadley-like’’ slower (about 10 m/s) overturning of the atmosphere from the equator to the high latitudes, with giant vortices at each pole recycling the air downward. No attempt to model the superrotation has been completely successful so far, indicating that the basic mechanisms of the phenomenon are unclear. More and better data, with improved spatial and temporal coverage, are needed in order to understand the detailed physical condition in this region and to modify the models to better reflect the reality. [14] Venus Express investigates the atmospheric dynamics by observing clouds at different levels [Drossart et al., 2007a; Markiewicz et al., 2007a], by deriving thermal winds from thermal profiles, and by monitoring airglow of different species. Specifically, it (1) measures the global wind fields in three dimensions and investigates whether the meridional circulation is one large basic ‘‘Hadley’’ cell extending from the surface to the upper atmosphere, or a variation of such a
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Figure 1. The abundance of the most important minor gases in the atmosphere of Venus as determined by missions before Venus Express. The individual gases are color coded as shown in the bottom left corner. The sizes of the blocks indicate the spread in several measurements rather than measurement errors. The bars with arrows indicate the limits of detection for the Venus Express instruments.
cell, or something completely different; (2) studies the temporal behavior of the polar vortices: how they evolve and how they couple to the two main components of the global circulation [Piccioni et al., 2007]; and (3) studies how the transition between the retrograde superrotation and the solar to antisolar circulation takes place in the interface between the mesosphere and the thermosphere. [Drossart et al., 2007b]. 2.3. Atmospheric Composition and Chemistry [15] The main components of the Venus atmosphere are CO2 (96.5%) and N2 (3.5%). Sulfur-bearing trace gases, carbon and chlorine compounds, and water vapor are present in the atmosphere in amounts from few to few hundred parts per million (ppm) [de Bergh et al., 2006]. Figure 1 shows the mean abundance of the main trace gases and their vertical profiles measured in the Venus atmosphere by earlier missions. [16] Although present in small amounts, the trace gases are involved in important complex chemical cycles. In the region around the cloud tops photochemical reactions between CO2, SO2, H2O, and chlorine compounds lead to the formation of sulfuric acid, which is the main component of the cloud particles. The chemistry of the lower atmosphere is dominated by thermal decomposition of sulfuric acid, and thermochemical cycles that include sulfur and carbon species and water vapor. Surface minerals can also play a significant role in buffering the abundance of certain gases in the lower atmosphere [Fegley et al., 1997]. The main sulfur-bearing gas, SO2, is present in the Venus atmosphere in amounts of a few hundred ppm, which is much more than expected from
thermal equilibrium with the surface minerals. Pioneer Venus measured a strong continuous decline of the SO2 abundance at the cloud tops during its 14 years of operation, indicating possible recent volcanic activity. [17] Venus Express measurements of the chemical composition of the atmosphere address the following: (1) The abundance and spatial and temporal variation of SO2, SO, H2O, HCl, and CO at the cloud tops, to improve the understanding of the physical and chemical processes in this region, including the production of sulfuric acid aerosols [Bertaux et al., 2007a], (2) vertical profiles of SO, SO2, H2O and HDO, HCL, and HF between 80 km and the cloud tops, and vertical profiles of CO from the cloud tops up to about 120 km, by stellar and solar occultation [Bertaux et al., 2007b], and (3) the abundance and spatial variation of H2O, SO2, COS, CO, H2O, HCl, and HF in the lower atmosphere, to improve the understanding of chemistry, dynamics, and radiative balance of the lower atmosphere, and to search for local volcanic activity [Drossart et al., 2007a; Svedhem et al., 2007b]. 2.4. Cloud Layer and Hazes [18] Venus appears completely featureless in visible light owing to a thick cloud layer located between 50 km and 70 km altitude. However, images made in the UV blue spectral range show much structure, both at large scale and at small scale (Figure 2). It has important implications for the energy balance since about half of the solar energy entering into the atmosphere is absorbed in this region. The nature of the UV absorbing matter remains one of the mysteries of Venus. Earlier observations have shown that the upper cloud layer consists of micron-sized droplets of 75% sulfuric acid.
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radiation penetrates through the atmosphere and heats the surface [Taylor et al., 2007]. The radiation balance, however, is not well constrained and in particular the spectral characteristics and the spatial and temporal variations of the outgoing radiation are not well known. [21] Venus Express measures the outgoing radiative flux (reflected solar light and thermal radiation from the atmosphere and the surface) over the near-IR spectral range and maps the temperature field and the cloud layer in three dimensions. This will constrain the optical models of the atmosphere and give an insight into the radiative and dynamical heat transport, and the role of the different species in the greenhouse mechanism and the planetary heat balance.
Figure 2. An image of Venus made by the Venus Monitoring Camera (VMC) instrument in ultraviolet light at 365 nm from a distance of about 35,000 km. The south pole is at the bottom right in the image and the equator at the top left. The cloud structure is shown to be very different in the polar region compared to the equatorial region. A bright elongated cloud is shown in the middle southern latitudes.
These droplets make up a major portion of the cloud mass but the processes and conditions for their formation are poorly understood, while the physical and chemical processes forming the rest of the cloud population are virtually unknown. [19] Venus Express sounds the structure, composition, dynamics, and variability of the cloud layer. Specifically, it (1) images the spatial distribution and temporal variability of the unknown UVabsorber [Markiewicz et al., 2007b; Drossart et al., 2007a]; (2) measures the vertical structure and microphysical properties of the upper cloud layer and the hazes above it, by stellar and solar occultation, limb observations, and nadir sounding [Bertaux et al., 2007a]; (3) measures the spatial variations of the opacity and particle sizes on the nightside (through the IR ‘‘spectral windows’’); this data together with information on the composition will help to constrain models of the cloud formation and evolution [Titov et al., 2006b]; and (4) searches for correlation in the spatial distribution of the unknown UV absorber with the abundance of SO2 and other trace gases. 2.5. Energy Balance and Greenhouse Effect [20] With a surface temperature of more than 730 K Venus has the strongest greenhouse effect in the solar system. The dense CO2 atmosphere is responsible for this extreme behavior, but water vapor and sulfur dioxide and the cloud layer also play a role [Crisp and Titov, 1997; Titov et al., 2007]. It is even more impressive when considering that Venus, in fact, absorbs less energy from the Sun than the Earth does, owing to the high albedo (76%), which is caused by the thick cloud cover of the planet. Less than 10% of the incoming solar
2.6. Plasma Environment and Escape Processes [22] Today Venus has only very little water, mostly in the form of water vapor. If it all were condensed it would form a global layer of about 3 cm, compared with 3 km on the Earth. There is no good reason why initially the two planets should have been significantly different and it is expected that Venus indeed has had large quantities of water, and possibly other volatile species, perhaps as much as the Earth. If so, how, when and why did this disappear? Pioneer Venus found that deuterium is enhanced relative to hydrogen about a factor 150 compared to the Earth’s value [Donahue et al., 1997]. This is an indication that hydrogen has been lost, while deuterium, having twice the mass, does not escape as easily as hydrogen. Therefore information on the present abundance of deuterium, hydrogen and oxygen and their escape rates are essential to the understanding of the history of water on Venus. [23] The lack of an internal magnetic field causes the solar wind to act on the upper atmosphere in a much more violent way than it does on the Earth. The solar wind interacts with the top of the ionosphere to form a complex system of plasma clouds, tail rays, filaments, and ionospheric holes on the nightside through which a substantial amount of material can leave the planet [Brace and Kliore, 1991]. The escape mechanisms induced by the solar wind are the dominant ones for the loss of heavy atmospheric gases such as oxygen because the gravitational force inhibits both Jeans escape and nonthermal escape. [24] To address the problems of atmospheric escape and investigate the plasma environment, Venus Express (1) determines the positions of the plasma boundaries for the different domains in the planetary environment, and their dependence on the solar activity [Zhang et al., 2008a, 2008b; Martinecz et al., 2008], (2) measures in situ, for the different domains, directional flux and energy of energetic neutral atoms, ions and electrons; of particular importance are the measurements of the escape rate of key species like hydrogen, deuterium and oxygen [Barabash et al., 2007a, 2007b], (3) measures the magnetic field in the different domains in the planetary environment [Zhang et al., 2006, 2007], (4) determines the vertical structure of the ionosphere, by radio occultation measurements [Ha¨usler et al., 2006; Pa¨tzold et al., 2007], and (5) measures the composition and energies of the undisturbed solar wind, as a reference for the Venus measurements and for comparison with similar measurements around the other planets of the solar system. The Venus Express measurements are taken at solar minimum activity, thus complementing the Pioneer Venus plasma studies that were acquired during solar maximum conditions.
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2.7. Surface Properties and Geology [25] Venus has a geologically young surface at an age of 700 Ma ± 200 Ma, as estimated on the basis of impact crater statistics. In addition there are indications that the age of the different regions do not differ much from each other. Volcanic and tectonic activities have strongly affected the surface [Solomon et al., 1992] forming highly deformed plateaus, tesserae, and extensive lowlands, planitiae. Many scientists favor a catastrophic global resurfacing of the crust, a mechanism unique among the terrestrial planets [Strom et al., 1994]. Important open questions are related to the surface mineralogy, the existence of present volcanic activity and the role of chemical interaction at the surface-atmosphere interface. Despite the fact that Venus Express is not carrying any instruments specifically devoted to surface science, it contributes to these studies in several ways [Titov et al., 2006b], by [26] 1. Investigating regions of high radar reflectivity, by bistatic measurements, focusing on highland regions like Aphrodite Terra, Beta and Atla Regio, and Maxwell Montes [Ha¨usler et al., 2006]. These areas have shown anomalously high reflectivity in the Magellan radar images. Unfortunately, these measurements have had to be discontinued since mid2007 owing to a still unexplained loss of power in the S band transmitter chain. [27] 2. Investigating the mass distribution in and around Atalanta Planitia by radio science gravity studies, through orbital trajectory analysis. [28] 3. Mapping the surface temperature and estimate the surface emissivity by observations in the 1 mm spectral window. [29] 4. Searching for surface hot spots and local deviations in chemical abundance in the lower atmosphere, in particular SO2, indicating possible volcanic activity. [30] 5. Searching for atmospheric waves generated by seismic activity and coupled to the atmosphere owing to the high atmospheric density at the surface [Drossart et al., 2007a].
3. Spacecraft [31] Venus Express uses the basic design of the Mars Express spacecraft, which was launched in 2003, adapted for the specific conditions at Venus [Sivac and Schirmann, 2009]. The main modifications to the spacecraft are related to the thermal control system and the power system. The most important characteristics are summarized in the following paragraphs. 3.1. Structure and Propulsion [32] The Venus Express spacecraft is based on a box-like structure with the dimensions 1.7 m 1.7 m 1.4 m. The distance from tip to tip of the deployed solar panels is about 8 m. The principal mechanical structures inside the spacecraft are the two fuel tanks, with MonoMethyl Hydrazine (MMH) as the fuel and Nitrogen TetrOxide (NTO) as the oxidizer, and a smaller helium tank for main tank pressurization. On both sides of these tanks two internal shear walls are built. The spacecraft, without solar panels, can be seen during the solar illumination testing in Figure 3 and the fully integrated spacecraft during its mating to the Fregat upper stage can
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be seen in Figure 4. The mass of the dry spacecraft is 633 kg, including 94 kg for the payload. The tanks were filled at launch with 570 kg of propellant, most of which was consumed during the orbit insertion maneuvers. Together with the launch vehicle adapter this made up the total launch mass of 1241 kg, leaving a comfortable margin to the maximum allowed mass of 1270 kg for the Soyuz Fregat launcher combination. Venus Express as well as Mars Express has been designed to sustain soft aerobraking, at a maximum dynamic pressure of 0.3 N/m2. 3.2. Thermal Control [33] The closer distance to the Sun, and the higher albedo of Venus compared to Mars, made it necessary to completely redesign the thermal control system. A new thermal blanket is used, giving the spacecraft a golden finish, in contrast to the black appearance of Mars Express. Several instruments and subsystems are connected to radiators mounted on different faces on the outside of the spacecraft. One side of the spacecraft, the -X side, carries the largest radiators which act as heat sinks for the coolers for the instruments requiring cryogenic temperatures. This side must never be exposed to the sun at any angle. The -Z side, where the main engine nozzle and the launch vehicle adapter interface reside, also must avoid illumination by the sun since these will absorb the solar heat very quickly. This design, together with specific operational constraints, ensures a low maximum temperature even in Venus orbit. Under some conditions the temperature would be too low, and therefore a set of electrical heaters are fitted to the most critical units. These heaters are either switched by mechanical thermostats or by the on board computer. 3.3. Power System [34] As the solar input in Venus orbit is about 2.6 kW/m2, which is about four times that of Mars or twice that of the Earth, the size of the solar panels could be significantly reduced. The original silicon-based solar cells have been replaced by gallium arsenide, optimized for operation at higher temperature. To keep the temperature in the proper range these cells are mounted in rows interleaved with optical solar reflectors, giving the solar panels a striped appearance. The electrical output in Venus orbit is approximately 1400 W. The remaining part of the power system, including the batteries, required only minor modification from the Mars Express version. 3.4. Telecommunication System [35] The spacecraft uses X band communications for both the telecommand uplink and the telemetry downlink. Fully redundant cross-strapped communication chains, including two 65 W traveling wave tube power amplifiers, feed the signal to the two high-gain antennas of 1.3 m and 0.3 m diameter. S band communications, with dual 5 W solid state power amplifiers, are included as a backup system and for use near the Earth during the first weeks following the launch. The data rate varies between 15 kbps to 228 kbps, corresponding to a downlink capability of between 400 Mbit to 6.5 Gbit per day, depending on the actual distance between Venus and the Earth. The antennas are body fixed and thus the spacecraft will turn toward the Earth for dumping the data
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Figure 3. The compete spacecraft excluding solar panels during the solar illumination test at Intespace, Toulouse, France. The facility was upgraded for this test in order to be able to produce the required beam with an intensity 2.6 kW/m2 over the full surface of the spacecraft. The front side shows the 1.3 m diameter high-gain antenna covered with its solar shield, a thin foil with germanium deposit to keep the antenna within its operational temperature during all spacecraft attitudes. The back side of the smaller high-gain antenna is the white dish seen on the top platform where all remote sensing instruments are also located. The top and front sides are allowed to be continuously illuminated by the Sun at Venus, while all other sides have strict limitations on the duration of solar illumination. The height of the structure is about 1.4 m.
each telemetry session, usually 8 – 10 h per day. The data are stored in a 12 Gbit solid state data recorder during the observation passes. [36] The radio science investigation makes full use of the telecommunication system and an ultrastable oscillator has been added to one chain of the system to enhance the performance for this purpose. The S band system experienced a drop in signal level during the solar conjunction in July 2006 and has therefore only been used occasionally since then. The reason for this anomaly is still not known. [37] The main ground station is the Cebreros station west of Madrid, Spain, which is used for all uplink and most data downlink activity. The New Norcia station in Australia is used for radio science activities. Occasionally the NASA DSN assists by giving extra coverage, for radio science activities in particular, but also in connection with periods of high data rate requirements, for example, when measurements at high temporal resolution are made, such as highresolution movies of the cloud motions. 3.5. Operational Constraints [38] While pointing to the Earth from Mars during communications the Sun is always within a cone of 40° from the Earth, making it easy to avoid illumination of protected surfaces. In addition the high-gain antenna acts a thermal shield of the spacecraft. From Venus, however, the Sun can appear anywhere in the sky, making it impossible to avoid
illumination of specific surfaces. To deal with this a second, smaller, high-gain antenna has been mounted on top of the +Z face of the spacecraft, pointing in the opposite direction from the main antenna. Then it is possible always to find an attitude where illumination of the forbidden faces can be avoided. Figure 5 shows that inside the quadrature, where the Venus-Sun-Earth angle is less than 45°, the small antenna is used; that is, the spacecraft needs to rotate 180° about the z axis at entry and exit of the quadrature. In addition the spacecraft will need to rotate 180° about the x axis at inferior and superior solar conjunction.
4. Payload [39] The payload is composed of seven instruments in three different categories; spectrometers and (spectral) imagers for remote sensing, plasma and magnetic field instruments for in situ measurements, and the ultrastable oscillator used for radio science. [40] Great care was taken in selecting the instrument complement in order to make optimum use of the limited resources on board. A complication was the short time available for development since the project schedule was very compressed. As it turned out, a very well balanced set of instruments could be assembled from existing designs and even from spare parts and units from other recent missions, mainly Mars Express and Rosetta. In addition, two completely new instruments
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Figure 4. The fully integrated spacecraft is being mounted on top of the Fregat upper stage. The black conical structure is the launch adapter, which remains on the Fregat stage after separation. Protective covers are still mounted on the folded solar panels and on the instrument apertures and the two cylindrical baffles of the star trackers. The white recessed surface on the front is the VIRTIS thermal radiator. were designed, developed and built in this short time, namely the VMC and the SOIR. [41] Figures 6 and 7 show how the different instruments complement each other with respect to their fields of view and their spectral range and resolution. Venus Express is the first dedicated atmospheric mission to Venus since the discovery of the near infrared spectral windows [Allen and Crawford, 1984; Baines et al., 2006]. The payload was composed in order to maximize the benefit from this new opportunity to study the atmosphere in three dimensions. Table 1 lists the individual instruments and their main functions and Figure 8 shows the locations of the instruments on the spacecraft. 4.1. Remote Sensing Instruments [42] Spectroscopy for the Investigation of the Characteristics of the Atmosphere of Venus/Solar Occultation in the Infrared (SPICAV/SOIR) [Bertaux et al., 2007a, 2009] is a set of three spectrometers optimized for providing thermal profiles and composition in the upper atmosphere by observations in stellar and solar occultation mode. It is also used in nadir mode, together with other instruments. The solar and stellar occultation technique provides very good vertical
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resolution and high sensitivity for species of low abundance (like SO2, COS, CO, HCL, HF, and others). Particularly important molecules to measure are H2O and HDO, since the ratio between the two carries information on the history of the total amount of water on the planet. Temperature and density profiles are retrieved by analyzing the absorption in various CO2 bands. The instrument covers three wavelength bands: 110 –310 nm, 0.7 –1.7 mm and 2.2 – 4.4 mm (SOIR). The SOIR unit utilizes a miniaturized Stirling-cycle engine to cool the detector to 90 K, and an acoustooptic tunable filter, together with a grating used at a high order of diffraction, to achieve a spectral resolution (l/dl) of more than 20,000. [43] The basic SPICAV unit has an important heritage from the SPICAM instrument on Mars Express, while the SOIR channel is a completely new development. Important recent findings include the determination of the D/H ratio as a function of altitude, the discovery of a warm layer on the nightside at the base of the thermosphere (90 km) and the discovery of a set of previously unobserved absorption lines of the CO2 isotope 16O12C18O [Bertaux et al., 2008; Wilquet et al., 2008]. [44] Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) is a combination of a visible (0.27– 1.1 mm) and an infrared (1.05– 5.2 mm) medium resolution imaging spectrometer (VIRTIS-M) with a high-resolution infrared (1.8 – 5.0 mm) spectrometer (VIRTIS-H) [Piccioni et al., 2009; Drossart et al., 2007a]. The infrared detectors are actively cooled to 80 K, while the visible sensor is passively cooled to below 200 K. The instrument is almost identical to the VIRTIS instrument presently flying on the Rosetta spacecraft. The imaging spectrometer has a field of view of 64 mrad 64 mrad, with a pixel size of 0.25 mrad 0.25 mrad. The images have a spatial dimension of 256 256 pixels and a spectral dimension of 432 lines. From near apocenter, the field of view covers about 1/3 of the diameter of Venus. Mosaics of 3 3 frames are constructed by repointing of the spacecraft in order to generate global images of the southern hemisphere. VIRTIS is addressing a large number of scientific questions and has generated dramatic images and video sequences of the south polar vortex, nonLTE emission patterns, profiles of abundance of several atmospheric gases, wind field maps and temperature profiles over the southern hemisphere, images of cloud structure at several altitudes and surface temperature and emissivity maps, to mention just a few. [45] Venus Monitoring Camera (VMC) is a small but efficient camera operating simultaneously in four narrow spectral bands at 365, 513, 965, and 1000 nm [Markiewicz et al., 2007a, 2009]. The UV band is mapping the cloud tops in reflected sunlight during daytime, where structure is visible owing to the still unknown UV absorber. These images are used for deriving global wind fields at this altitude and for studying the cloud morphology. The near-IR 1mm band is used for mapping the surface brightness during nighttime. The remaining two bands include O2 airglow emission and possible water vapor absorption. VMC has an unusual design with four optical chains, each one with its own permanent filter, sharing one CCD, with each optical system having its own dedicated quadrant on the CCD. When the CCD is read out in normal mode each image frame actually contains four images. The field of view is 17.5° (0.3 rad), resulting in a pixel footprint size on the surface ranging from 200 m at
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Figure 5. The orbit of Venus in a Sun-Earth fixed frame, with the Sun at the center and the Earth at the left. The attitude of the spacecraft while Earth pointing will change depending on the relative position in this frame in a way such that the Sun is illuminating only two of the sides of the spacecraft (+X and +Z). To achieve this, the spacecraft is rotated 180° around the z axis at the points marked quadrature and 180° around the x axis at inferior and superior conjunction.
pericenter to 45 km at apocenter. VMC has sampled a large number of images, and important findings include the determination of global wind fields at different times, the characterization of the cloud morphology in different regions and synthesis of surface maps. [46] Planetary Fourier Spectrometer (PFS) is a high spectral resolution double pendulum spectrometer with two channels, a short wave channel (0.9 – 5.5 mm) and a long wave channel (5.5 – 45 mm) [Formisano et al., 2006, 2009]. The history of this instrument goes back to the Russian Mars-96 mission that unfortunately was lost in a launch failure in 1996. It flew again on Mars Express where it performs very well and has made important discoveries. Some of the main objectives for Venus were determination of the temperature field in the altitude range 55– 100 km, both on the nightside and on the dayside, mapping of the surface temperature, and making profiles of the abundance of a large number of gases in the middle and lower atmosphere. It was also to measure the outgoing thermal flux to determine the global radiation budget. Unfortunately, a mechanism that controls the mirror that switches the beam between the different calibration targets and the view toward the planet is stuck in its launch position, pointing toward a blackbody target. In spite of many attempts it has not been possible to move the scanner since the launch. 4.2. In Situ Instruments [47] Analyzer of Space Plasma and Energetic Atoms (ASPERA-4) is the fourth of this series of instruments to
fly in space, but the first of its kind around Venus [Barabash et al., 2007a, 2009]. ASPERA-4 is identical to ASPERA-3 presently in orbit around Mars on Mars Express. The instrument has four different sensors, housed in two separate units. The ion mass analyzer, IMA, measures ions separated by mass up to m/q = 40 for energies from 10 eV/q to 36 keV/q. The ions are mapped at 22.5° resolution over 360° in azimuth and at 4.5° resolution over ±45° in elevation with respect to the instrument orientation. The Neutral Particle Imager (NPI) measures the integral flux of neutral particles between 100 eV and 60 keV with an instantaneous field of view of 9° by 344° at an angular resolution of 4.6° by 11.5°. The Neutral Particle Detector (NPD) consists of two identical units that are basically pinhole cameras for the energy range 0.1 – 10 keV, each with a field of view of 9° by 90° and a resolution of 5° by 30° with separation of hydrogen and oxygen. The Electron Spectrometer (ELS) is a compact electrostatic analyzer for the energy range 1 eV to 15 keV with an energy resolution of 7%. The field of view is 10° in elevation and 360° in azimuth with a resolution of 22.5° in azimuth. ASPERA has made a large number of measurements in all accessible domains around Venus and in the solar wind and has characterized the different regions and their boundaries with respect to flux, energy and composition of the detected particles. A major result is the absolute determination of the escape rates of H+, O+ and He+ which is important for estimating the historic water content on the planet. [48] Magnetometer (MAG) is a dual sensor fluxgate instrument with the main sensor mounted on the tip of a 1 m
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Figure 6. A comparison of the fields of view (FOV) of the remote sensing instruments onboard and the size of Venus corresponding to a distance of 20,000 km. The background image was taken by VMC at ultraviolet wavelengths from about 35,000 km. The south pole is at the bottom of the image. At apocenter the planetary disk is smaller than the VMC FOV and about three times the VIRTIS-M FOV.
long boom and an auxiliary sensor mounted at the foot of the boom on the outer surface of the spacecraft [Zhang et al., 2006, 2009]. This dual sensor configuration has proven very useful since the spacecraft was developed without a specific magnetic cleanliness program and therefore disturbances from the spacecraft itself are unavoidable. The data from the auxiliary sensor are used to clean the data of the main sensor, and a performance similar to that of a magnetically clean spacecraft is achieved. Venus Express is the first spacecraft where this concept has been used. MAG is the only instrument on board that is making continuous measurements. Samples are taken at 1 Hz frequency when the spacecraft is in the solar wind and at up to 128 Hz around the pericenter. The results include detailed models of the bow shock and the induced magnetopause, and detection of foreshock activities and upstream waves. Particularly interesting is the detection of whistler waves that are interpreted as evidence of lightning in the atmosphere. 4.3. Radio Science [49] Venus Radio science (VeRa) is the radio science investigation [Ha¨usler et al., 2006, 2009]. The main objective is
to determine the temperature and the density of the lower atmosphere in the altitude range 40– 90 km and to determine the electron density of the ionosphere up to the ionopause by means of radio occultation of the telemetry signal. Additional objectives are to carry out bistatic radar measurements over selected areas of the surface, including those highly reflective (at radar wavelength) areas discovered by the Magellan radar, in particular in areas of high elevation, like Maxwell Montes and Thetis Regio. Measurements are made by directing the main S/C antenna toward the area of interest on the surface, and receiving the reflected signal on a large antenna back on Earth. This was attempted successfully several times until the power in the S band channel was lost, as described above. The absorption in the atmosphere at S band is manageable but at X band it is too high for reliable operation. [50] VeRa occultation profiles have been used to map the thermal structure in the lower atmosphere at a high vertical resolution and medium horizontal resolution, separating the daytime and the nighttime regions. The lower parts show little difference between day and night but in the higher regions variations are clearly visible.
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Figure 7. A comparison between the wavelength ranges and spectral resolutions of the Venus Express remote sensing instruments.
Figure 8. The Venus Express spacecraft in a semitransparent view, showing the positions of the seven scientific instruments. The top face, the +Z platform, accommodates all the apertures of the remote sensing instruments. This side is turned toward Venus during observations. The two ASPERA units are mounted on the bottom side ( Z platform). 10 of 19
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Table 1. The Seven Instruments of the Venus Express Payloada Name
Function
SPICAV/SOIR
UV-IR spectrometer
VIRTIS VMC PFS ASPERA-4 MAG VeRa
Measured Parameters
Major Topics of Interest
Wavelength 110 – 310 nm, Solar and stellar occultation. Upper atmosphere: 0.7 – 1.65 mm, 2.2 – 4.4 mm Composition, D/H ratio, thermal profiles. Resolving power up to 20,000 UV-Vis-IR Virtis-M, 0.27 – 5.2 mm imaging Atmospheric dynamics, structure and composition. Imaging spectrometer Virtis-H 1.8 –5.0 mm single point at resolving power 2000 Surface temperature, emissivity. 4 band camera Wavelength 365 nm, 513 nm, 965 nm, 1000 nm Wind speed, clouds, surface, global context Fourier spectrometer Wavelength 0.9 – 45mm, resolving power 1200 Thermal profiles, composition. Presently non operational Energetic particles Electrons 1eV – 15 keV, ions 0.01 – 36 keV/q, Atmosphere solar wind interaction, neutral particles 0.1 – 60 keV atmospheric escape. Magnetic fields B field, 8 pT – 262 nT, at 128 Hz Induced magnetosphere. Plasma boundaries. Lightning Occultation of telecom link X- and S-band, Doppler shift, Lower atmosphere density and thermal structure. polarization, amplitude Ionosphere. Surface radar reflectivity.
a The total mass is 92 kg. The instruments together produce between 1 Gbit and 6 Gbit data per day, adapted to the telemetry transfer rate which depends on the actual Venus-Earth distance.
[51] The radio measurements also contribute to surface science by precise tracking of the spacecraft trajectory, from which data the gravity field is determined and gravity anomalies are derived. The gravity field of Venus is in general well known from previous missions, but a few anomalies need to be further investigated, for example in the region of Atalanta Planitia. A secondary objective is to study the solar corona at times near Venus superior conjunction, when the signal from Venus passes close to the disk of the Sun, revealing the structure, density and dynamics of the solar corona.
[54] Many scientific and technical constraints were considered when the operational orbit was selected. To achieve good global coverage, a polar obit was deemed to be essential. The need for a global view and long duration measure-
5. Mission Scenario and Operational Orbit [52] Venus Express was launched by a Soyuz-Fregat combination from Baikonur, Kazakhstan, on 9 November 2005 and arrived at Venus on 11 April 2006 (Figure 9). The injection into the heliocentric transfer orbit by the launcher was very precise and very little fuel had to be spent for trajectory corrections during the cruise to Venus. During the first part of the 153 days cruise all spacecraft and payload elements were checked out for proper operation. [53] The Venus orbit insertion was divided into several steps, starting with a capture maneuver which included a 52 min long burn by the 400 N main engine, to achieve the required delta-v of 1251 m/s [Warhaut and Accomazzo, 2009]. This delta-v is significantly higher than that required for the capture of Mars Express into Mars orbit. The reason is twofold; the differential velocity between the spacecraft and the planet is larger at Venus, and the larger mass of Venus requires more energy to reduce the apocenter height once captured. The insertion burn placed the spacecraft in a 9-day orbit with an apocenter height of 330,000 km. During this initial ‘‘capture’’ orbit, six blocks of scientific observations were included to benefit from the unique opportunity of having a very high apocenter distance that allowed global observations and dynamic studies on a large scale. After finishing the capture orbit, a sequence of smaller engine burns took place to reduce the apocenter height to the operational value. During the capture maneuver and the subsequent orbit reduction burns a total of 482 kg of fuel was consumed. The final operational orbit was reached on 6 May 2006.
Figure 9. Venus Express liftoff on a Soyuz-Fregat launcher from Baikonur Cosmodrome, Kazakhstan, 0333 UT, 9 November 2005. The spacecraft will subsequently be inserted into a perfect heliocentric trajectory toward Venus.
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ments for dynamic studies, together with close-up studies and near-planet in situ measurements, drove the requirement toward a highly elliptical orbit. A low circular orbit was not desired but would also not be feasible owing to the high fuel demands. An orbit with a pericenter height of 250 km and an apocenter height of 66,000 km, resulting in a period of 24 h, was selected. Such an orbit also had the advantage of allowing operators work at the same time each day and, since the main ground station is located in Spain, this turns out to be normal daytime working hours, saving a significant amount in the operational costs. [55] The remaining parameter to decide was the latitude of the pericenter. Since Venus rotates very slowly, only one revolution in 243 Earth days, there is very little polar flattening and consequently very little drift in the pericenter latitude, only about 3° per year. This makes the decision more critical for Venus than for Mars or the Earth. The orbital trajectory at arrival left basically only two options free for the choice of pericenter latitude, either 15° or +78°. In order to study one hemisphere at large distance and the other at short distance +78° was selected. This choice also nicely complements Pioneer Venus, which had its pericenter around the equator. The final orbit is shown in Figure 10. [56] A specific feature with a highly elliptical, polar orbit around Venus is the drift in pericenter altitude, due to perturbations by the solar gravity, which causes the pericenter altitude to drift upward or downward at a rate of 1 – 3 km per day. For the case of Venus Express the drift is upward until May 2009, then downward. This drift is compensated for by regular thrusting by four 10 N thrusters, to maintain the pericenter height between 250 km and 400 km and, after August 2008, between 175 km and 275 km. This will ultimately lead to the depletion of the fuel and an end of the mission in late 2013 if no countermeasures are taken. A possible alternative is to use aerobraking in the upper atmosphere of Venus to reduce the apocenter height (and the orbital period), as well as the rate of the drift of the pericenter height. A 12 h (37,000 km apocenter height) or 8 h (27,000 km apocenter height) orbit would dramatically reduce the drift rate and allow operation of the mission for another decade if required.
6. Science Planning and Operations [57] The planning of the scientific operations is a complex process that is carried out at different levels in a sequential manner. At the highest level, governed by the Science Requirements Documents, the Venus Express Science Working Team develops a long-term Science Activity Plan (SAP) typically covering a year or more of observations [Titov et al., 2006b]. This plan identifies specific orbital, and other, characteristics like solar illumination, accessible surface targets, solar eclipse periods, and Earth occultation periods. A scheme for the operations is then designed, on the basis of 10 different building blocks, called Science Cases, which can be put together in different combinations. [58] A medium-term plan (MTP) covering a period of 4 weeks, is agreed between representatives of the instrument teams on the basis of the guidelines in the SAP, approximately 3 months before its execution. [59] The lowest level in the planning cycle is the short-term plan (STP) which is prepared on a weekly basis. Here the
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Figure 10. The operational 24 h orbit around Venus. There are four major operational phases: (1) Around apocenter, at 10 h to 14 h, full disk imaging is made by VIRTIS and VMC, (2) along the ascending branch, between 14 h and 22 h, longduration observations are made for studies of the dynamics and cloud evolution, and (3) around pericenter, 23 h to 1 h, high spatial resolution observations and occultation measurements are performed. (4) Data transmission to the Earth takes place with the spacecraft Earth pointing from 2 h to 10 h orbital time. detailed command files with all settings for the instruments are included. The whole process is based on an exchange of a large number of files of different complexity between the PI teams, the science operations center (VSOC) and the mission operations center (VMOC), following a well-defined scheme [Koschny et al., 2009]. [60] The orbit is normally split into four parts where different activities take place (Figure 10). The region around apocenter is used for global mapping by VMC and synthesis of 3 3 frame mosaics for full disk coverage by VIRTIS. The ascending branch, from 10 h before pericenter to about 2 h before pericenter, has a long duration view of the same area of the planet and is used for studies of atmospheric dynamics. At times of high data rate, time lapse movies are composed here. The pericenter region, between about 1 h before and 1 h after pericenter, is shared between high spatial resolution
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Figure 11. The present coverage of the VIRTIS-M images of the southern hemisphere as a function of latitude and local solar time is shown as a color-coded image. The coverage of the latitudes between 50°S and the south pole is obviously very good, while the coverage at the equator and northward of the equator is very limited.
nadir measurements and stellar occultation measurements. If in a solar eclipse period or an Earth radio occultation period, solar and radio occultation measurements also take place here. The descending branch of the orbit, from 2 h after pericenter to apocenter, is normally reserved for telecommunication and data downlink.
7. Coverage of Observations [61] A global and systematic survey of the atmospheric and plasma phenomena is a major goal of the Venus Express mission. To separate spatial variations from temporal variations, good coverage in a multiparameter space is essential for most measurements. Therefore priority is given to areas with poor coverage as far as possible. In general the coverage of the southern hemisphere is good and the coverage of the northern hemisphere is limited. This is inherent in the present orbit that has a larger distance to the planet, allowing more time for measurements over the southern hemisphere. [62] The coverage of VIRTIS spectral images over the southern hemisphere is shown in Figure 11. It can be seen that the polar region is well covered and the nightside of the southern hemisphere is reasonable well covered south of 50° southern latitude. On the day side, the equatorial region and the midlatitudes have very poor coverage. In addition the coverage of the northern hemisphere is very poor. Figure 12 shows the VIRTIS surface coverage from measurements dur-
ing nighttime. These data are essential for making surface maps. It can be seen that the coverage at midlatitudes between longitudes 150° and 30° is fairly good. The rest of the southern hemisphere is less well covered, in particular the longitudes between +45° to +150°. Also the surface coverage in the northern hemisphere is very poor. [63] The VMC dayside images provide uniform coverage of the southern hemisphere. These observations are used for cloud morphology and wind tracking studies. In the north only small-scale imaging is possible owing to the close proximity of the planet and the fast motion of the spacecraft. The VMC surface observations are limited to low latitudes (±40°). These images are made during nighttime, but owing to stray light only images taken during eclipse are useful. The observations complement well VIRTIS thermal mapping of the southern hemisphere. However, full coverage of all longitudes at low latitudes would require about 7 years since the part of the surface seen in eclipse drifts slowly because the rotation period of the planet is close to a Venus year. [64] The coverage of the SPICAV stellar occultation profiles depends on the availability of sufficiently bright UV stars. Presently, about 30 stars are used by SPICAV UV for occultation studies of the upper atmosphere, covering latitude ranging from 50°S to 40°N. A given star will always occult Venus at the same latitude but at varying local time, resulting in a coverage appearing as horizontal lines (Figure 13). The SPICAV and SOIR solar occultations by definition always
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Figure 12. The present coverage of the VIRTIS-M nighttime images of the southern hemisphere as a function of latitude and longitude is shown as a color-coded image. The coverage of the latitudes between 40°S and 80°S for longitudes between 150° and 30° is fairly good, while the coverage of other locations is limited. These data are required for seeing the lower atmosphere and the surface. occur at the terminator, i.e., 0600 or 1800 local solar time (Figure 14b), except very close to the poles owing to the slight tilt of the axis of rotation of the planet. Full coverage in local solar time is thus not possible, although improved temporal coverage, to study both short-term and long-term variations, can be achieved by more measurements. [65] Figure 14a shows the coverage of the VeRa Earth radio occultations. It can be seen that even after three full occultation seasons the coverage is not very dense, in particular in the northern hemisphere, from the equator up to about 60° northern latitude. [66] ASPERA has a very good coverage inside the bow shock, except for in a small corridor straight south of the planet. It also has a good set of reference data taken in the undisturbed solar wind, well outside the bow shock (Figure 15). [67] MAG is the only instrument that is almost constantly on. In the solar wind it samples at a fixed frequency of 1Hz and close to pericenter it samples at 128Hz for a few minutes. This thorough coverage provides data for a multitude of processes within and outside the induced magnetosphere.
8. Mission Extension Topics [68] Presently, the Venus Express mission is funded for operations until May 2009. Most of the original primary objectives will have been met by then, but a number of old and new questions that have arisen would benefit from more data and longer total mission duration. As described above, the coverage for many parameters both in local solar time and
in latitude and longitude is not yet sufficient for detailed analysis. The present solar minimum has surprised us by being longer than normal, and the solar activity is expected to increase soon. It is very valuable to monitor the response by the atmosphere to this increase and to compare it to the situation with the low solar activity we have had over the past 2 years, and to the Pioneer Venus measurements at high solar activity in the years around 1980. An improved knowledge on the influence of the solar activity on the escape rate of various species may have important implications for the understanding of the evolution of the atmosphere. [69] Long-duration monitoring of key species, for example SO2, and of cloud properties and wind fields is essential to make conclusions on secular variation of composition and dynamics. Likewise, long-duration observations of the surface increase the likelihood of finding local hot spots and/or volcanic activity. An extended operational life will also allow time for a further reduction of the pericenter altitude, perhaps to 170 km in order to improve the basis for the in situ measurements of energetic particles and magnetic fields. The thermal structure and density in the otherwise inaccessible region between 170 and 190 km can be studied by atmospheric drag measurements. [70] An extension beyond 2010 would open up the possibility for joint observations with the Japanese Planet-C spacecraft. Planet-C is focusing on small-scale atmospheric dynamics and weather phenomena [Nakamura et al., 2007], while Venus Express aims at a more general study of the atmosphere. Planet-C will be placed in a highly elliptical
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Figure 13. Coverage of the SPICAV stellar occultations in the ultraviolet channel; latitude versus local time. From these measurements, thermal profiles in the altitude range 90 km to 140 km are derived together with and abundances of several atmospheric gases. The coverage is poor at high latitudes, and more data will be needed for a full coverage. These measurements are only effective over the dark hemisphere.
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equatorial orbit, complementing well the polar orbit of Venus Express. Joint observations would allow simultaneous highresolution imaging and global context imaging, and detailed 3-D imaging of the morphology of the upper clouds. This would be of particular interest for the studies of cloud dynamics and cloud evolution. [71] As discussed above an attractive way to extend the lifetime of the mission is to significantly lower the apocenter altitude by aerobraking. This would allow both an extended duration of the present observations and allow new investigations from the vantage of a different orbit. In addition density and temperature of the upper atmosphere can be retrieved from the drag exerted on the spacecraft by the atmosphere, parameters that in this altitude (140– 170 km) range cannot easily be measured remotely. Aerobraking in 2012 is under consideration and will be studied in detail in the near future.
9. Venus Express and the ESA Solar System Science Program
Figure 14. (a) Coverage of VeRa radio occultation data; latitude versus local time. From these measurements, profiles of atmospheric density and temperature are derived in the altitude range 35 km to 90 km. Additional data will be needed for a full coverage. (b) Coverage of the SOIR solar occultation measurements. By principle, only data along the terminator are available. More data will be needed for coverage at high southern latitudes and between 40°N and 80°N.
[72] Since Venus Express is very similar to Mars Express, with several instruments in common and simultaneous operations, interesting comparisons can be made. Together with data from Earth orbiting science satellites like Envisat and the Cluster spacecraft, an important basis for the comparative planetology of the terrestrial planets with atmospheres is laid. Possibilities also open up for simultaneous investigations of the solar wind and atmosphere interactions, and ‘‘space weather’’ studies on the three planets. The Solar System Science program of ESA is addressing all major bodies of the inner solar system from the Sun to Mars, and asteroids and comets. The three planets mentioned above all have ESA spacecraft presently orbiting them and the Sun is being observed, jointly with NASA, by the SOHO spacecraft from L1. [73] The last of the terrestrial planets, Mercury, will be studied in depth, jointly with Japan, by the dual spacecraft mission BepiColombo, which will be launched in 2014. The Rosetta spacecraft is already on its way to a comet and is due to arrive at its target, comet 67P/Churyumov-Gerasimenko, in 2015. During the cruise two asteroid flybys are scheduled. Thus, a fairly complete portion of the inner solar system is being researched, and synergistic effects can also be expected even if not predictable in advance. The information gathered will form a basis for the definition of other missions still to come. [74] ESA’s science program for the future, ‘‘Cosmic Vision,’’ with launches foreseen in the timeframe 2015 – 2025, is addressing a number of basic scientific themes with questions of a fundamental nature, like ‘What are the conditions for life and planetary formation?’ and ‘How does the Solar System work?’. A large number of missions were proposed for Cosmic Vision and several of them are now being studied in a competitive phase for a final selection in 2009. As discussed above, Venus Express already contributes to parts of these objectives, in particular in the field of comparative planetology and planetary evolution. [75] The ESA Planetary Science Archive, PSA, stores the data from all missions and makes them available to the world wide community. The first data from Venus Express, as well
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Figure 15. The ASPERA coverage shown in Venus solar cylindrical coordinates (positive x axis pointing toward the Sun and the y axis showing the distance toward the Venus-Sun line). The coverage is excellent in most of the volume inside the bow shock, with a slightly weak coverage only straight south of the southern hemisphere. A very good set of solar wind reference data are available from measurements around the apocenter. as the other missions, are available for immediate electronic download at www.rssd.esa.int/psa.
10. Conclusions [76] Originally born as a ‘‘mission of opportunity’’ reusing the Mars Express design and available instruments, Venus Express has proven to be the most powerful mission ever in Venus orbit. A fortunate combination of a versatile spacecraft, a state of the art payload, and an efficient ground segment allows scientists to carry out a complex and systematic survey of the planet from the surface to the thermosphere and above. The observations include nadir and limb geometry, stellar, solar and Earth occultation measurements, and in situ plasma investigations. These studies have unveiled details of the atmospheric structure, composition, cloud morphology, dynamics and escape processes never observed before. [77] The papers following this one in this special section of JGR report on the most important findings to date. Many of the original objectives have been addressed to a significant
depth, while additional objectives are being formulated for the extended mission presently being planned. The analysis of the large amounts of data collected is a formidable task, and the science teams are engaged with both operations and processing the data. There is much more to be done with the data sets already acquired and new types of observations still to be carried out, such as atmospheric drag measurements and coordinated measurements with Planet-C, are promising new exciting results. [78] Everyone involved in the mission is appreciating the worldwide revival of the interest in Venus science that Venus Express appears to have triggered. The spacecraft and the instruments are in a good condition, and should continue to provide new data for scientists worldwide for several years to come. [79] Acknowledgments. The authors are grateful for the professionalism and enthusiastic commitment shown by all colleagues in the experimental teams, at EADS-Astrium, ESTEC, ESOC, and ESAC, which contributed greatly to the positive spirit and the success of the Venus Express mission.
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Lellouch, E., T. Clancy, D. Crisp, A. J. Kliore, D. Titov, and S. W. Bougher (1997), Monitoring of mesospheric structure and dynamics, in Venus-II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher et al., pp. 295 – 324, Univ. of Ariz. Press, Tucson. Markiewicz, W. J., et al. (2007a), Venus Monitoring Camera for Venus Express, Planet. Space Sci., 55, 1701 – 1711, doi:10.1016/j.pss.2007. 01.004. Markiewicz, W. J., et al. (2007b), Morphology and dynamics of the upper cloud layer of Venus, Nature, 450, 633 – 636, doi:10.1038/nature06320. Markiewicz, W. J., et al. (2009), VMC: The Venus Monitoring Camera, in Venus Express, edited by A. Wilson, Eur. Space Agency Spec. Publ., ESA SP-1295, in press. Martinecz, C., et al. (2008), Location of the bow shock and ion composition boundaries at Venus: Initial determinations from Venus Express ASPERA-4, Planet. Space Sci., 56, 780 – 784, doi:10.1016/j.pss.2007.07.007. Moroz, V. I. (2002), Studies of the atmosphere of Venus by means of spacecraft: Solved and unsolved problems, Adv. Space Res., 29, 215 – 225, doi:10.1016/S0273-1177(01)00571-3. Nakamura, M., et al. (2007), Planet-C: Venus climate orbiter mission of Japan, Planet. Space Sci., 55, 1831 – 1842, doi:10.1016/j.pss.2007.01.009. Pa¨tzold, M., et al. (2007), The structure of Venus’ middle atmosphere and ionosphere, Nature, 450, 657 – 660, doi:10.1038/nature06239. Piccioni, G., et al. (2007), South polar features on Venus similar to those near the north pole, Nature, 450, 637 – 640, doi:10.1038/nature06209. Piccioni, G., et al. (2009), VIRTIS: The Visible and Infrared Thermal Imaging Spectrometer, in Venus Express, edited by A. Wilson, Eur. Space Agency Spec. Publ., ESA SP-1295, in press. Seiff, A., J. T. Schofield, A. J. Kliore, F. W. Taylor, S. S. Limaye, H. E. Revercomb, L. A. Sromovsky, V. V. Kerzhanovich, V. I. Moroz, and M. Y. Marov (1985), Models of the structure of the atmosphere of Venus from the surface to 100 km altitude, Adv. Space Res., 5, 3 – 58, doi:10.1016/ 0273-1177(85)90197-8. Sivac, P., and T. Schirmann (2009), The Venus Express spacecraft system design, in Venus Express, edited by A. Wilson, Eur. Space Agency Spec. Publ., ESA SP-1295, in press. Solomon, S., et al. (1992), Venus tectonics: An overview of Magellan observations, J. Geophys. Res., 97, 13,199 – 13,255. Strom, R. G., G. G. Schaber, and D. D. Dawson (1994), The global resurfacing of Venus, J. Geophys. Res., 99, 10,899 – 10,926, doi:10.1029/ 94JE00388. Svedhem, H., et al. (2007a), Venus Express - The first European mission to Venus, Planet. Space Sci., 55, 1636 – 1652, doi:10.1016/j.pss.2007. 01.013. Svedhem, H., D. V. Titov, F. W. Taylor, and O. Witasse (2007b), Venus as a more Earth-like planet, Nature, 450, 629 – 632, doi:10.1038/nature06432. Svedhem, H., et al. (2009), Venus Express: Mission overview, in Venus Express, edited by A. Wilson, Eur. Space Agency Spec. Publ., ESA SP1295, in press. Taylor, F. W. (2006), Venus before Venus Express, Planet. Space Sci., 54, 1249 – 1262, doi:10.1016/j.pss.2006.04.031. Taylor, F. W., et al. (1980), Structure and meteorology of the middle atmosphere of Venus: Infrared remote sounding from the Pioneer Orbiter, J. Geophys. Res., 85, 7963 – 8006, doi:10.1029/JA085iA13p07963. Taylor, F. W., H. Svedhem, and D. Titov (2007), Venus Express and terrestrial planet climatology, in Exploring Venus as Terrestrial Planet, edited by L. W. Esposito, E. R. Stofan, and Th. E. Cravens, Geophys. Monogr. Ser., vol. 176, pp. 157 – 170, AGU, Washington D. C. Titov, D. V., E. Lellouch, F.W. Taylor, L. Marinangeli, H. Opgenoorth, and Venus Express Team (2001), Venus Express: An orbiter for the study of the atmosphere, the plasma environment, and the surface of Venus, Mission Definition Rep. ESA-SCI (2001)6, Eur. Space Agency, Paris. Titov, D. V., H. Svedhem, and F. W. Taylor (2006a), The atmosphere of Venus: Current knowledge and future investigations, in Solar System Update, edited by P. Blondel and J. W. Mason, pp. 87 – 110, Springer, Berlin. Titov, D. V., et al. (2006b), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017. Titov, D. V., M. Bullock, D. Crisp, N. Renno, F. W. Taylor, and L. V. Zasova (2007), Radiation in the atmosphere of Venus, in Exploring Venus as Terrestrial Planet, edited by L. W. Esposito, E. R. Stofan, and Th. E. Cravens, Geophys. Monogr. Ser., vol. 176, pp. 121 – 138, AGU, Washington D. C. Warhaut, M., and A. Accomazzo (2009), The ground segment and mission operations, in Venus Express, edited by A. Wilson, Eur. Space Agency Spec. Publ., ESA SP-1295, in press. Wilquet, V., et al. (2008), Line parameters for the 01111-00001 band of C12O16O18 from SOIR measurements of the Venus atmosphere, J. Quant. Spectrosc. Radiat. Transfer, 109, 895 – 905, doi:10.1016/j.jqsrt. 2007.12.021.
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Zhang, T. L., et al. (2006), Magnetic field investigation of the Venus plasma environment: Expected new results from Venus Express, Planet. Space Sci., 54, 1336 – 1343, doi:10.1016/j.pss.2006.04.018. Zhang, T., et al. (2007), Little or no solar wind enters Venus’ atmosphere at solar minimum, Nature, 450, 654 – 656, doi:10.1038/nature06026. Zhang, T. L., et al. (2008a), Initial Venus Express magnetic field observations of the magnetic barrier at solar minimum, Planet. Space Sci., 56, 790 – 795, doi:10.1016/j.pss.2007.10.013. Zhang, T. L., et al. (2008b), Initial Venus Express magnetic field observations of the Venus bow shock location at solar minimum, Planet. Space Sci., 56, 785 – 789, doi:10.1016/j.pss.2007.09.012.
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Zhang, T. L., et al. (2009), MAG: The fluxgate magnetometer of Venus Express, in Venus Express, edited by A. Wilson, Eur. Space Agency Spec. Publ., ESA SP-1295, in press. H. Svedhem and O. Witasse, ESA, ESTEC, Keplerlaan 1, NL-2200 AG Noordwijk, Netherlands. (
[email protected]) F. Taylor, Department of Atmospheric, Oceanic, and Planetary Physics, Oxford University, Oxford OX1 3PU, UK. D. Titov, Max-Planck Institute for Solar System Research, Max-PlanckStrasse 2, D-37191 Katlenburg-Lindau, Germany.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, E00B07, doi:10.1029/2008JE003074, 2008
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A latitudinal survey of CO, OCS, H2O, and SO2 in the lower atmosphere of Venus: Spectroscopic studies using VIRTIS-H E. Marcq,1,2 B. Be´zard,1 P. Drossart,1 G. Piccioni,3 J. M. Reess,1 and F. Henry1 Received 14 January 2008; revised 8 April 2008; accepted 10 July 2008; published 23 September 2008.
[1] The high-resolution channel (R ’ 2000) of the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) instrument (VIRTIS-H) aboard Venus Express has provided numerous spectra of the nightside infrared thermal emission in the 2.3-mm window. Mixing ratios of various minor species in the 30–40 km range could therefore be inferred using this spectral window at higher latitudes accessible to the spacecraft but which cannot be observed from Earth. The previously known enhancement in carbon monoxide (CO) toward high latitudes is confirmed and extended up to 60° with a mixing ratio varying from 24 ± 3 to 31 ± 2 ppmv at 36 km. Measurements of carbonyl sulfide (OCS) also agree with the previously suspected latitudinal variations that are anticorrelated with those of CO, ranging between 2.5 ± 1 and 4 ± 1 ppmv at 33 km. New constraints were also derived on the mean abundance of water vapor (H2O, 31 ± 2 ppmv) and sulfur dioxide (SO2, 130 ± 50 ppmv) in the probed altitude range. CO and OCS variations are interpreted as caused by large-scale vertical motions, an explanation under current testing by various chemical and dynamical modeling. In such a case, these variations may help constrain the chemical time scale of those species in the lower troposphere. Citation: Marcq, E., B. Be´zard, P. Drossart, G. Piccioni, J. M. Reess, and F. Henry (2008), A latitudinal survey of CO, OCS, H2O, and SO2 in the lower atmosphere of Venus: Spectroscopic studies using VIRTIS-H, J. Geophys. Res., 113, E00B07, doi:10.1029/2008JE003074.
1. Introduction [2] The discovery of Venus’ nightside near infrared emission in 1983 by Allen and Crawford [1984] provided a new, valuable method to investigate the deep atmosphere of the planet. The physical origin of this emission was rapidly understood in the 1980s, on the basis of its physical properties and successful modeling [Crisp et al., 1989; Kamp et al., 1988]. The hot and deep layers of the lower atmosphere emit strong thermal radiation that can partly escape the thick atmosphere in a few narrow spectral windows thanks to the sub-Lorentzian behavior of CO2 line shape and the weakly absorbing scattering by the H2SO4 particles within the overlying cloud layers. In the 2.3-mm window considered hereafter, the emission originates from the 30– 40 km altitude range (average temperature of about 500 K). [3] Besides the imaging of the lower cloud deck whose opacity modulates the intensity of these emissions [Carlson 1 Laboratoire d’E´tudes Spatiales et d’Instrumentation en Astrophysique, UMR8109, CNRS, Observatoire de Paris, Meudon, France. 2 Laboratoire de Me´te´orologie Dynamique, UMR8539, CNRS, Universite´ Paris 6, Paris, France. 3 Istituto di Astrofisica Spaziale e Fisica Cosmica, Istituto Nazionale di Astrofisica, Rome, Italy.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003074$09.00
et al., 1991], the main scientific interest comes from its spectroscopic analysis. The comparison with radiative transfer synthetic models [Be´zard et al., 1990; Kamp and Taylor, 1990; Pollack et al., 1993; de Bergh et al., 1995] provided numerous constraints on the abundance of minor gaseous constituents. In the 2.3-mm window, carbon monoxide (CO), carbonyl sulfide (OCS), water vapor (H2O, HDO), sulfur dioxide (SO2) and hydrofluoric acid (HF) were detected and measured thanks to their spectral signatures. Furthermore, the spectral behaviors of CO and OCS are rich enough so that not only their mean abundance level, but also their respective vertical gradients could be derived, and were shown to be opposite by Pollack et al. [1993], with CO increasing with altitude as it is produced by the photochemistry of CO2 above 60 km. [4] The study of local compositional variations began later, in the 1990s, when Collard et al. [1993] noticed in the NIMS/Galileo spectra recorded during the flyby that the spatial variations of the CO band shape indicated a northern enrichment at higher latitudes of about 35%. More recently, spectra have been acquired at the NASA IRTF using the spectrometer SpeX during the quadratures that occurred since 2003. Their better spectral resolution compared to NIMS allowed us to confirm the CO northern enrichment and report a similar enrichment in the southern hemisphere, as well as a correlated depletion of OCS [Marcq et al., 2005, 2006] in the portions of Venus’ nightside that can be
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Table 1. Summary of Mentioned Science Cases Science Case
Description
1 2 3 5 7
pericenter nadir observations off-pericenter nadir observations apocenter mosaic observations stellar occultations limb observations
observed from Earth (below 40° in latitude). A latitudinal variability of the vertical gradient of these two species in the probed layers could also be measured thanks to the high signal-to-noise ratio in the best data sample, and it appears that both gradients were steeper at higher latitudes than in the equatorial region [Marcq et al., 2006]. Spatial variations in water vapor have also been sought by several observers [Drossart et al., 1993] since the localized enrichment reported by Bell et al. [1991], which was hypothesized to be due to cloud subsidence and evaporation. [ 5 ] Keeping in mind that CO increases and OCS decreases with increasing altitude, the latitudinal variability of CO and OCS can be qualitatively understood as being primarily caused by the global meridional dynamics. At high latitudes, downward fluxes from the Hadley cell circulation bring air depleted in OCS and enriched in CO, whereas the ascending branch of these cells has the opposite effect close to the equator. Quantitative modeling is currently progressing (Y. Yung et al., Modeling the distribution of OCS in the lower atmosphere of Venus, submitted to Journal of Geophysical Research, 2008; E. Marcq et al., Simulations of the latitudinal variations of CO and OCS below the clouds of Venus using a general circulation model, submitted to Journal of Geophysical Research, 2008). Besides the use of minor species as tracers of the deep circulation, the study of the compositional variability can also be related to the geological activity of the surface; a local enrichment in water vapor and SO2 and a correlated modification of the D/H ratio of water may indicate an ongoing volcanic eruption [Donahue et al., 1997; Grinspoon, 1993; Donahue, 1999]. [6] After more than a year of activity, the VIRTIS instrument aboard ESA’s Venus Express spacecraft has provided numerous spectra from the nightside of Venus very similar to the SpeX/IRTF ones. Using very similar methods to our previous studies, we present in this paper the new constraints we have derived from these data about the spatial variations of CO, OCS and water vapor, as well as a semiquantitative interpretation of the aforementioned variations.
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probe, there is an asymmetry in the observation times for both hemispheres: the northern one is seen from a nearer vantage point, but for shorter durations. [8] Within each 24-hour orbit, 8 hours are dedicated to communication with the Earth. The remaining 16 hours are used by the six functioning instruments (ASPERA, MAG, SPICAV-SOIR, VeRa, VIRTIS, VMC) for the various Venus-related observations (such as limb or nadir sounding, solar or stellar occultations) according to preplanned schedules named ‘‘science cases.’’ A brief summary of the science cases mentioned in this article is available in Table 1 [from Titov et al., 2006b]. 2.1.2. VIRTIS-H [9] The VIRTIS spectral imager is described in detail by Coradini et al. [1998]. Here we describe the characteristics of the H-channel that are relevant to our analysis. The specific optical device of the H-channel consists of an echelle grating spectrometer, and dispersion orders from 8 to 15 are focused on the detector plane. The 2.3-mm emission can be seen on three among these orders, designated as orders 5, 6 and 7. The spectral resolution of each of these spectra varies with the wavelengths, as shown on Figure 1. The spectra are then sampled on 432 discrete wavelengths for each dispersion order, although the first tens of pixels on the right side (short wavelengths) of the focal plane array usually yield smaller radiances than expected because of the low grating efficiency on the edges of the orders, thus reducing their effective spectral ranges. Another peculiarity of the array reading system that we had to take into account in our processing methods is the difference in fluxes between odd and even pixels along a given spectrum. Nevertheless, this peculiarity is partially corrected by the calibration process, yielding a few percent difference in the absolute intensities between even and odd pixels. 2.2. Quality of the Observations 2.2.1. Orbits [10] We have currently reviewed all VIRTIS-H spectra between orbits 7 (27 April 2006) and 298 (13 February 2007). The proportion of orbits during which nightside spectra could be successfully acquired is close to 40%. This
2. Observations 2.1. VIRTIS and the Venus Express Mission 2.1.1. Brief Overview of the Mission [7] A detailed description of the Venus Express mission is provided by Titov et al. [2006a]. Here we summarize the mission characteristics relevant to this study. The orbit of the spacecraft is quasi-polar, with a periapsis located at 78°N, and highly elliptical (the distance to the surface ranging from 250 km up to 66000 km) with a 24-hour period. As a consequence of the Keplerian motion of the
Figure 1. Spectral resolution versus wavelength for orders 5 (red), 6 (green), and 7 (blue).
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Table 2. Processed Orbits Orbit
Number of Processed Spectra
Science Case
Latitude on Venus
Single Spectrum S/N Ratio
007 024 027 033 040 042 060 061 073 074 076 077 078 079 082 088 095 096 097 098 100 102 103 104 110 110 111 113 116 125 128 129 134 136 139 143 148 243 248 249 258 277 284 298
128 100 35 32 50 40 160 140 230 230 128 192 300 200 80 80 50 1728 40 1664 30 30 30 50 2880 150 640 50 50 32 1400 320 1200 200 40 96 50 160 1100 150 250 96 51 200
N/Aa 1 1 7 7 7 7 2 1 5 1 2 1 1 1 7 2 2 1 2 2 2 5 2 2 2 2 2 2 5 2 2 2 2 2 2 2 1 3 1 5 1 2 1
29°S – 26°S 20°S – 55°N 20°N – 60°N 20°N – 50°N 0° – 60°N 40°N – 60°N 10°S – 70°N 30°S – 50°N 30°S – 20°N 30°S – 20°N 35°N – 55°N 44°S – 40°S 10°S – 70°N 0° – 70°N 25°N – 60°N 30°N – 70°N 40°N – 70°N 58°S – 44°S 40°N – 60°N 55°S – 28°S 45°N – 55°N 30°N – 55°N 30°N – 50°N 10°N – 55°N 50°S – 15°S 10°N – 60°N 49°S – 24°S 35°N – 60°N 35°N – 60°N 35°N – 60°N 60°S – 45°S 11°S – 50°N 60°S – 45°S 56°S – 54°S 20°N – 70°N 10°S – 5°N 11°N – 25°N 25°S – 40°N 43°S – 40°S 0° – 60°N 30°S – 0° 20°S – 5°S 20°S – 50°N 10°S – 40°N
10 3 – 10 8 – 12 4–8 4–6 6 3–8 4–8 4 – 12 10 – 30 4–8 5 2–4 4 – 25 2–6 4 – 18 4–8 3–4 3 3–4 4 5 5 2.5 4 4 2–8 3 3 3 4 4 4 10 3–4 3 5 12 10 20 6 20 4 – 10 4 – 14
a
During insertion orbit.
proportion and the temporal distribution of these orbits are consistent with the expected requirements (VIRTIS operational and pointing to a location on the nightside). Nevertheless, a subset of these orbits (14% of the total number) show a S/N ratio significantly better than those of the typical spectra taken at similar latitudes. They are summarized in Table 2. The retrievals shown in section 3.2 have therefore been conducted on this subset. 2.2.2. Influence of the Observing Mode (‘‘Science Case’’) [11] Quite surprisingly, the periapsis observations (case 1) were not the ones which consistently yielded the best spectra. It can be understood when we consider that the limiting factor in the processing of the VIRTIS-H spectra is the signal-to-noise ratio. In order to improve it, we have to integrate signal for as long as possible: at least a couple of minutes (around 20 frames in case 1) and up to half an hour, and a few hundred spectra for off-pericenter (case 2) or apoapsis mosaic observations (case 3), when the orbital velocity is slower than for periapsis observations. Longer
integration times are not suitable for a study of spatial variations, since the location on Venus of the various spectra would be blurred owing to the motion of the spacecraft. As a comment, it should be noted that even if it could theoretically be done, it would be useless to shorten the integration time to enhance the spatial resolution below roughly 30 km, since multiple scattering of the thermal emission by the cloud particles causes an intrinsic blurring of the signal over such a horizontal scale. Actually, our integration signal causes a latitudinal extension of the spectra’s location over a few degrees at most (longitudinal extension can be neglected owing to the quasi-polar orbital motion), thus enabling an integrated S/N ratio around 100 (individual spectra’s S/N ratios are given in Table 2). [12] Considering this trade-off between spatial resolution and signal-to-noise ratio, it appears that case 1 observations are well suited to low-latitude regions, where the signal is strong enough to cope with the moderate integration durations available in this science case owing to the high orbital velocity. On the other hand, midlatitudinal to high-latitudinal
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Figure 2. Superimposition of averaged nightside spectra of the same location on orders 5 (red), 6 (green), and 7 (blue) and their respective zero level. Note also the difference of spectral resolution between the various orders, in agreement with Figure 1. regions require a slower orbital velocity which can ensure a longer integration in spite of an increased distance. This explains why high-latitude regions have been mostly successfully observed during science cases 2, 5 or even 3. 2.3. Comparison with Similar Earth-Based Spectra [13] Sets of very similar spectra in terms of spectral range (1.92 – 2.52 mm) and resolution (R = 2000) have been acquired since 2003 using SpeX/IRTF by Marcq et al. [2005]. These observation campaigns were intended as a benchmark for the methods used hereafter, so that a comparison between these previous studies [Marcq et al., 2005, 2006] and the present one is very interesting. However there are some differences between both data sets that should be highlighted. 2.3.1. Spatial and Temporal Coverage [14] First of all, the spatial and temporal coverage is of course much better for VIRTIS, a dedicated instrument aboard a space probe orbiting Venus, than for the groundbased observations. From Earth, the nightside of Venus can be successfully observed only for a few weeks between inferior conjunction and maximal elongation. This period of favorable observing geometry occurs only twice in every 19.2 months. In contrast, VIRTIS has been providing spectra of the nightside of Venus on a regular basis (40% of the orbits on average) since April 2004 and will continue until at least the end of the extended mission in 2009. Another interesting feature of the Venus Express observations is that the quasi-polar orbit enables the observation of northern and southern circumpolar regions which are otherwise unreachable; the farthest observations with respect to the Venusian equator using SpeX could only reach 40 degrees of latitude. 2.3.2. Preliminary Processing [15] Another major advantage of space-based observations is the absence of any sunlight reflected from the dayside in virtually all nightside spectra, thanks to the apparent size of Venus as seen from the spacecraft. For SpeX observations, the much smaller angular diameter of
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Venus and the fact that a significant portion of the visible Venusian disk is illuminated near maximal elongation can lead to a strong scattering of the much brighter light from the solar illuminated portion of the planet onto the nightside, reducing the attainable S/N. The removal of the diffused component is a significant step in the processing and a major source of uncertainty for the retrievals of Marcq et al. [2005] and Marcq et al. [2006]: it prevented the observation of the nightside hemisphere close to the terminator, thus contributing to the narrowness of the latitudinal range of Earth-based observations. [16] Nevertheless, some processing is also needed for VIRTIS-H spectra, since the 2.3 mm window is covered by 3 different orders of dispersion that should agree on the actual intensity observed within their common spectral ranges. A typical superimposition is shown on Figure 2 for orders 5, 6 and 7. The differences on the actual value of the zero-intensity level are clearly visible between order 6 on one hand and orders 5 and 7 on the other hand. In the subsequent processing, the actual zero level for order 6 is computed from the zero levels measured on the orders 5 and 7, whose mutual agreement despite being on opposite sides of the 2.3-mm window make more trustworthy than order 6’s short-wavelength zero level. [17] As a last comment, Earth-based observations must also be corrected for the spectral signature of the terrestrial atmosphere, especially for water vapor which is much more abundant on Earth than on Venus. Such a concern is absent for VIRTIS observations, which makes them more reliable, especially for water vapor retrievals. 2.3.3. Signal-to-Noise Ratio [18] One major asset of Earth-based observations compared to VIRTIS-H ones is their quality. The longer integration time and especially the much greater light collecting area available more than compensate for the much greater distance between the instrument and Venus. A typical individual VIRTIS-H spectrum in a low-latitude to midlatitude region has a S/N ratio of between 5 and 10. Averaging over a few tens of spectra improves this ratio to about 50, at the expense of a having a spatial precision roughly equivalent to that of SpeX observations with a seeing of about 1 arcsec. SpeX spectra which are comparable in terms of integration time, location on Venus and latitudinal extension show an S/N ratio of 100 or higher. Thus, rather than supplementing Earth-based observations, VIRTIS provides a long-needed complementary spatial and temporal coverage of nightside deep atmosphere.
3. Retrievals 3.1. Algorithms 3.1.1. Synthetic Model [19] The forward synthetic model that we use is exactly the same as described by Marcq et al. [2006]. A typical fit is shown on Figure 3. [20] A major flaw of all radiative transfer models used to investigate the 2.3-mm window is the poor fit in the CO2 dominated spectral domain between 2.2 and 2.3 mm. Therefore, we have tried the new CDSD-750 database for CO2, suited for high-temperature calculations as encountered in the probed altitude region. Unfortunately, the fit we obtained was worse compared to the same model using the
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Figure 3. Comparison in the 2.2– 2.5 mm interval of a good quality order 6 spectrum (from orbit 277, in black) with the synthetic model. The relatively poor agreement at wavelengths shortward of 2.3 mm is due to the poor knowledge of CO2 opacity. older HITEMP database, whether a continuum absorption for CO2 was considered or not. So we discarded the CDSD-750 database, reverted to the HITEMP database and kept the continuum absorption of CO2 to the usual constant value of 3.5 108 cm1.amagat2 throughout the window [Pollack et al., 1993; Marcq et al., 2006], owing to the lack of experimental measurements at the relevant pressures and temperatures. [21] An update of the temperature profile that we consider in all of our simulations has also been tried from the older mean VIRA profile [Seiff, 1983] to the newer VIRA-2 [Moroz and Zasova, 1997]. A comparison between the two profiles in the altitude range where the emission originates from (30 to 40 km) showed very little difference, especially for the vertical gradient dT/dz. This latter could have been a concern since altering dT/dz distorts the simulated spectra in almost the same way as a change in CO abundance; it affects the relative brightness temperature of neighboring atmospheric layers where CO mixing ratio differs. On the contrary, changing the mean temperature of the 30– 40 km region while maintaining the temperature gradient produces almost no shape distortion of the spectra, but merely alters the integrated flux in the whole window, as expected. As far as we are concerned, the differences were too weak to justify a change of our temperature profile. [22] Still, the assumption of a horizontally uniform temperature profile could be invalid in the polar regions (above 60° in latitude), where some variability of T(z) is expected even in the lower atmosphere [Pa¨tzold et al., 2007]. But these regions yield a too weak signal in VIRTIS-H to be studied, so that we maintained this assumption. It is also worth noting that the weaker fluxes of the polar region are caused by a thicker lower cloud cover, whose modal distribution of aerosols is poorly known. A proper modeling of nightside polar spectra would therefore require a good understanding of both temperature profile and vertical
and modal distribution of cloud particles, which are both difficult to assess yet. [23] Finally, the synthetic spectra were sampled and convolved according to the instrumental function to allow straightforward comparison between synthetic and observed spectra. The instrumental functions differ for orders 5 and 6, and their width and shape vary with wavelength in a complex way. 3.1.2. One-Parameter Retrieval [24] The algorithm is based on a simple calculation of the residuals between an observed spectrum and various sets of precomputed synthetic spectra of various prescribed atmospheric parameters (e.g., vertical profiles of minor constituents, cloud opacity) within a few specific spectral intervals. This algorithm is almost exactly similar to the one described in section 4.1.1. by Marcq et al. [2006], except for some minor adaptations due to the differences between VIRTIS-H and SpeX. [25] Actually, besides the different determination of the true zero level (see 2.3.2) and, more generally, the completely different preprocessing, especially regarding diffused sunlight removal for SpeX, the key differences come from the fact that VIRTIS-H nightside spectra are available on two dispersion orders (5 and 6; order 7 was only used for zero-level determination since it does not cover the minor species absorption bands). Furthermore, the totality of the 2.3-mm window is only available on order 6, as can be seen on Figure 2: the 2.2-2.3 mm interval is not accurately rendered in order 5. From these considerations, the following adaptations were required: [26] 1. The former version of the algorithm used the same spectral interval to compute the free-scaling parameter ai (designed to have the best possible agreement between a given synthetic spectrum Si and the observed spectrum E) and to compute the residuals between E and Si. Owing to the lack of coverage for order 5, the algorithm is now able to use spectral intervals that are distinct for these two uses, the
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Figure 4. Spectral intervals used by our algorithm for the various species. The thick lines correspond to the residual computation intervals, whereas the whole lines (thin and thick) correspond to the scaling parameter determination intervals.
scaling interval typically being a superset of the residual estimation interval. The actual values, which can differ from order to order, are shown on Figure 4. [27] 2. Orders 5 and 6 of VIRTIS-H overlap between 2.3 and 2.5 mm. Although the analysis on both orders is done separately since they yield two distinct spectra of the same location, the estimates are considered as two independent measurements of the same spectrum: the resulting estimation e and its 2 uncertainty s are given by the following s6 e5 þs25 e6 5 s6 ffi , where the subscripts formulæ: e = s2 þs2 and s = psffiffiffiffiffiffiffiffiffi 2 2 6
5
s5 þs6
5 and 6 refer to the considered order. To ensure the consistency of such a combination, the retrievals were dismissed if their maximal normalized probability density, 2 5Þ ], was inferior to a threshold depending given by exp[ðes62e þs25 6 on the species, usually between 0.8 and 0.9. Note that coadding spectra from orders 5 and 6 would not make sense since their spectral resolution differ on their common wavelength range. 3.1.3. Strategy for Retrieving Gaseous Profiles [28] The global retrieval strategy is mostly unchanged since Marcq et al. [2006], where the 2.3-mm window was studied sequentially from the shorter wavelengths to the longer ones. The spectral intervals used for the retrievals of these three species are shown on Figure 4. Note that wavelengths less than 2.35 mm could not be used for order 5-based retrievals, since the instrument underestimates the intensity at these wavelengths for this order. 3.1.3.1. Carbon Monoxide Retrieval [29] The CO absorption band centered at 2.33 mm cannot be properly observed on order 5, so that only order 6 can be used for the retrievals. Spectral distortions due to the variation of the lower cloud opacity between the various spectra are taken into account as by Marcq et al. [2006] by first measuring the cloud opacity using integrated spectral radiance between 2.2 and 2.5 mm and then accordingly interpolating the synthetic spectra to the measured opacity. The considered set of CO vertical profiles are all proportional to the standard VIRA-2 profile. 3.1.3.2. Water Vapor Retrieval [30] Spectra from both orders 5 and 6 can be used for this species. Variations of lower cloud opacity, which affect spectral radiance in the whole window, and carbon monoxide, measured at the previous step and whose absorption band overlaps the water absorption band, are taken into account in the same way as previously stated. The set of synthetic profiles are here also proportional to the standard VIRA-2
profile, with no attempt to explore various D/H ratios at this stage. 3.1.3.3. Carbonyl Sulfide Retrieval [31] Spectra from orders 5 and 6 were also processed for the retrievals. Spectral overlapping with water vapor and carbon monoxide requires that the variations of these two species are taken into account, not to mention the consideration of cloud opacity variations here also. The set of considered synthetic profiles are various vertical translations of the standard VIRA-2 profile. 3.1.3.4. Sulfur Dioxide Retrieval [32] Its spectral signature being very narrow, sulfur dioxide retrievals are performed tentatively with a very low accuracy, using orders 5 and 6. Variations of carbonyl sulfide, water vapor and lower cloud opacity are taken into account, with the SO2 profiles assumed to be vertically uniform. 3.2. Results 3.2.1. CO [33] The mixing ratio of carbon monoxide at 36 km inferred at various locations is plotted on Figure 5. The minimal value of 24 ± 3 ppmv appears to be slightly off the equator, between 0° and 20°S. Its abundance then rises symmetrically with respect to this minimum toward higher latitudes, reaching 31 ± 2 ppmv at 60°S and between 30 and
Figure 5. Variations of the CO abundance at 36 km according to various VIRTIS-H spectra (orbits shown in various colors). Error bars stand for 1-s uncertainty levels.
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3.2.3. H2O [35] As concerns water vapor, no similar latitudinal trend could be observed on our results shown on Figure 7. Nevertheless, the combination of order 5 and 6 retrievals has enabled us to enhance the accuracy on its mean value significantly: 31 ± 2 ppmv between 30 and 40 km. Some temporal and/or longitudinal variability may also be possible with slightly drier conditions (29 ± 1 ppm) past 20°N, but this departure with the mean value is not significant: we have no evidence for any variations within this data set. 3.2.4. SO2 [36] Sulfur dioxide measurements could not reach a high accuracy considering the narrowness of its spectral band on the long-wavelength edge of the 2.3-mm spectral window, so that no latitudinal variation could be detected (see Figure 8). The assumption that the mixing ratio is uniform yields a mean mixing ratio of 130 ± 50 ppmv. Figure 6. Variations of the OCS abundance at 33 km according to various VIRTIS-H spectra (orbits shown in various colors). Error bars stand for 1-s uncertainty levels. 40 ppm at 60°N, where the coverage is not dense enough to ensure as much accuracy as in its southern counterpart. The dispersion of the measurements suggests some longitudinal and/or temporal variability within less than the 300-day interval during which the processed spectra have been acquired, although never exceeding 20% of the zonal mean value nor the pole-to-equator gradient. 3.2.2. OCS [34] Carbonyl sulfide latitudinal variations are shown on Figure 6. These variations appear to be anticorrelated to those of carbon monoxide, showing a maximal value of 4 ± 1 ppmv between 20°S and the equator, which was also the latitude where CO was minimal; the asymmetry of CO and OCS variations with respect to the equator could be related to the presence of high-elevation landmasses (namely, Aphrodite Terra, which is centered around 10° S) through the perturbations they cause to the general circulation. The minimum value of 2.5 ± 1 ppmv is reached near 60°S. The relative uncertainty on these retrievals is quite high, higher than what the S/N value or the accuracy of the fitting could account for, and such that we cannot dismiss a latitudeindependent mixing ratio between 3 and 3.5 ppm. Our current interpretation of this poor precision is related to the choice of the 1-D vertical synthetic profile set mentioned in section 3.1.3, which implies a univocal relation between mean abundance and vertical gradient. This relation prevents the exploration of the whole 2-D parameter space to which spectra are sensitive (i.e., vertical gradient and, independently, mean abundance in the 30– 36 km range). The explored subset is probably out of the optimal fit region since the dispersion of measurements is consistently inferior to the 1-s error bars, which points out that statistical noise is not the only source of discrepancy between synthetic and observed spectra. The same concern was raised by Marcq et al. [2006], and was solved by enabling the vertical gradient and the abundance to be retrieved separately. Such a study requires very high S/N ratio, and is therefore still in progress for the best VIRTIS-H spectra currently available.
3.3. Comparison with other Spectroscopic Studies 3.3.1. Lower-Resolution Spectra [37] Observations at lower spectral resolution could constrain with a similar accuracy the species with the most prominent band in the whole 2.3-mm window, namely, CO. H2O and OCS retrievals are currently in progress at a lesser precision compared to VIRTIS-H retrievals. Nevertheless, these studies require less signal to work on, so their horizontal coverage is substantially larger. The observations of Collard et al. [1993] using NIMS/Galileo were the first of this kind, and pointed to a northern enrichment in CO of 35 ± 15% from the equator to 47°N, in a good agreement with ours (30 ± 15% in the same latitudinal range). More recently, Tsang et al. [2008] has used VIRTIS-M spectra to map the CO abundance over large portions of the night hemisphere, and their conclusions about the latitudinal trends for CO corroborate ours in both hemispheres: from 23 ± 2 ppm near the equator to 32 ± 2 ppm at 35 km around 60° –70°. 3.3.2. Similar Resolution or Higher-Resolution Spectra [38] Spectra at a resolution similar or higher than ours in the 2.3-mm spectral window have been acquired and pro-
Figure 7. Variations of the water vapor at 33 km according to various VIRTIS-H spectra (orbits shown in various colors). Error bars stand for 1-s uncertainty levels.
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Figure 8. Variations of the sulfur dioxide at 35 km according to various VIRTIS-H spectra (orbits shown in various colors). Error bars stand for 1-s uncertainty levels. cessed before, but focused first on obtaining horizontally averaged spectra rather than studying their spatial variations. Comparisons are nevertheless useful, and agree with our results. The mean values of the CO mixing ratio at 36 km given by Pollack et al. [1993], Be´zard et al. [1990] and Taylor et al. [1997] range between 20 and 30 ppm, consistent with our own measurements. Using the same spectra, OCS and SO2 were also constrained, and respectively found to be 4.4 ± 1 ppm at 33 km (OCS [Pollack et al., 1993]) and 130 ± 40 ppm between 30 and 40 km (SO2 [Be´zard et al., 1993]). These determinations are also in full agreement with our own averages. The mean H2O mixing ratio between 30 and 40 km was found using high-resolution (R 20000) CFHT spectra in the 1.74 and 2.3-mm windows by de Bergh et al. [1995] to be 30 ± 10 ppm, which is also fully compatible with ours. In contrast, the lower estimate of 26 ± 4 ppm found by Marcq et al. [2006] is possibly due to contamination by terrestrial water vapor: the mean abundance is probably closer to 30 ppmv than to 26. [39] More recently, latitudinal studies in this spectral window were conducted using SpeX/IRTF spectra [Marcq et al., 2005, 2006] in order to prepare the processing of VIRTIS-H spectra. These results also agree with ours, at least in the latitude interval that can be observed from Earth: CO was found to vary between 26.5 ± 0.5 ppm (36 km) near 40°S and 23.75 ± 0.75 ppm near the equator, which matches our results. OCS variations were similarly anticorrelated with CO, although quantitative comparisons are harder considering the large vertical gradient of OCS in this region since their results are given at a different altitude of reference (36 km instead of 33). 3.4. Dynamical Interpretations 3.4.1. Qualitative Interpretation [40] The current interpretation of the anticorrelated variations of CO and OCS relies upon our knowledge of the global-scale circulation, especially vertical motions. It is known [Pollack et al., 1993; Marcq et al., 2006] that in the 30– 40 km altitude range, CO mixing ratio increases with height whereas OCS decreases. These observed vertical gradients are also predicted by most chemical models
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[Krasnopolsky, 2007; Krasnopolsky and Pollack, 1994]: CO is produced by photo-dissociation of CO2 by UV radiation above the clouds and consumed in the lower atmosphere as a reducing agent, whereas OCS is produced at ground level by reactions involving pyrite, CO and CO2 whereas it is strongly oxidized by SO3 (coming from the evaporation of H2SO4 at the lower cloud floor). The Hadley-cell vertical circulation of Venus causes upward motion in the equatorial region (due to solar heating) and downward motion at higher latitudes; most global circulation models (GCMs) predict the peak of downward velocity around 60° (S. Lebonnois et al., Superrotation of Venus’ atmosphere analysed with a full General Circulation Model, manuscript in preparation, 2008). Therefore, there is a tendency to deplete the CO mixing ratio around the equator (and increase that of OCS, which is more abundant at lower levels) whereas the trends are opposite at higher latitudes, which is precisely the observed pattern for CO and OCS. We can also note that the other measured species (H2O and SO2), which do not exhibit latitudinal variations, have no detectable vertical gradient either. 3.4.2. Estimate of Typical Relaxation Times [41] If the explanation given above is true, we can estimate the order of magnitude of the chemical relaxation time toward thermochemical equilibrium profiles for CO and OCS, assuming that the latitudinal contrasts we observe are caused only by the large-scale vertical motions. This coarse approach neglects other phenomena such as diffusion, and is very simplistic as for the chemical modeling, so it cannot yield more than order-of-magnitude estimates of the chemical time scale. [42] The advected net flux of a tracer whose mixing ratio ~ (q~ is q is given in a 1-D vertical model by r v) = w dq dz dw v stands for the wind speed vector, w for the dz q, where ~ vertical wind velocity and z for the altitude. Neglecting the variation of w between 30 and 40 km, this flux is in balance 0 ð zÞ where t with the chemical relaxation @q@tðzÞjchem = qðzÞq t stands for the characteristic time and q0(z) the expected vertical profile in the absence of any dynamical perturbation. In a stationary regime, this yields t¼
q0 q w dq dz
q q0 is typically around a few ppm, jdq/dzj about 1 ppm/ km for CO and OCS [Pollack et al., 1993; Taylor et al., 1997; Marcq et al., 2006] and GCMs predict typical velocities below 1 mm/s. This gives for both species characteristic times t 106 – 107, s in order to allow the formation of the equator-to-pole gradients in CO and OCS seen here. We would like to stress that this value is not related to a definite chemical process, but should merely be seen as a threshold: a faster chemistry (with a smaller t) would not allow enough time for the vertical circulation to establish the latitudinal contrasts described in section 3.2. Other characteristic times could be expressed regarding other processes (e.g., horizontal circulation, eddy diffusion), but a more detailed description is beyond the scope of this paper. 3.4.3. Quantitative Studies in Progress [43] To progress further in the interpretation of CO and OCS variations, a number of modeling efforts are currently
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in progress. The one from Yung et al. (submitted manuscript, 2008) focuses on OCS and adopts an accurate chemical scheme embedded in a 2-D circulation model, whereas Marcq et al. (manuscript in preparation, 2008) models both CO and OCS-type tracers with a simple relaxation scheme within a full 3-D radiatively consistent GCM. The results of these two models can be found in the above-quoted papers.
4. Conclusion [44] First of all, these results stress the complementary role of VIRTIS-M and VIRTIS-H channels of Venus Express, the precision of abundance retrievals and the number of species being the main assets of VIRTIS-H in this field whereas the CO mapping capabilities of VIRTIS-M over large extents on the Venus nightside are not accessible here. We have also significantly improved the accuracy of water vapor retrievals thanks to the availability of two spectral orders over the 2.3-mm spectral window. Even using a single order, CO variations were confirmed and validated up to 60° in latitude for both hemispheres, still fully compatible with previous estimates by Collard et al. [1993] and Marcq et al. [2005] (30 ± 15%). Anticorrelated OCS variations first noticed from Earth were also tentatively confirmed and extended in the same latitudinal range, though the relatively poor accuracy of the retrievals cannot rule out a constant value; the accuracy of the OCS retrievals should greatly benefit from a more detailed study with OCS abundance and vertical gradient being derived separately as it was the case for Marcq et al. [2006]. [45] With the oncoming extended phase of Venus Express, the prospective of future observations are very good. Exploration of high-latitude region (beyond 60°) could be possible thanks to the implementation of new science cases (enabling the tracking of a specific location during longer times), although such a study would probably require new additional information (better knowledge of possible horizontal variations of the T(p) profile, modal and vertical distribution of aerosol particles). Supportive studies from space and/or Earth in the other spectral windows could also help in constraining the vertical profiles of several minor species at different altitudes, especially H2O. [46] The first results of coupled dynamical-chemical simulations are also expected to improve significantly our understanding of the chemistry below the clouds of Venus, especially regarding the relative time scales of various dynamical and chemical phenomena. Observational constraints provided by the present paper are also of paramount interest for developing and improving such complex coupled simulations. [47] Acknowledgments. This work has been supported and funded by the Centre National d’E´tudes Spatiales (CNES).
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B. Be´zard, P. Drossart, F. Henry, E. Marcq, and J. M. Reess, Laboratoire d’E´tudes Spatiales et d’Instrumentation en Astrophysique, UMR8109, CNRS, Observatoire de Paris, 5 place Juleps Janssen, F-92195 Meudon, France. (
[email protected]) G. Piccioni, Istituto di Astrofisica Spaziale e Fisica Cosmica, Istituto Nazionale di Astrofisica, Via Fosso del Cavaliere 100, I-00133 Roma, Italy.
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Tropospheric carbon monoxide concentrations and variability on Venus from Venus Express/VIRTIS-M observations Constantine C. C. Tsang,1 Patrick G. J. Irwin,1 Colin F. Wilson,1 Fredric W. Taylor,1 Chris Lee,2 Remco de Kok,1 Pierre Drossart,3 Giuseppe Piccioni,4 Bruno Bezard,3 and Simon Calcutt1 Received 26 January 2008; revised 10 June 2008; accepted 25 June 2008; published 1 October 2008.
[1] We present nightside observations of tropospheric carbon monoxide in
the southern hemisphere near the 35 km height level, the first from Venus Express/Visible and Infrared Thermal Imaging Spectrometer (VIRTIS)-M-IR. VIRTIS-M data from 2.18 to 2.50 mm, with a spectral resolution of 10 nm, were used in the analysis. Spectra were binned, with widths ranging from 5 to 30 spatial pixels, to increase the signal-to-noise ratio, while at the same time reducing the total number of retrievals required for complete spatial coverage. We calculate the mean abundance for carbon monoxide at the equator to be 23 ± 2 ppm. The CO concentration increases toward the poles, peaking at a latitude of approximately 60°S, with a mean value of 32 ± 2 ppm. This 40% equator-to-pole increase is consistent with the values found by Collard et al. (1993) from Galileo/NIMS observations. Observations suggest an overturning in this CO gradient past 60°S, declining to abundances seen in the midlatitudes. Zonal variability in this peak value has also been measured, varying on the order of 10% (3 ppm) at different longitudes on a latitude circle. The zonal variability of the CO abundance has possible implications for the lifetime of CO and its dynamics in the troposphere. This work has definitively established a distribution of tropospheric CO, which is consistent with a Hadley cell circulation, and placed limits on the latitudinal extent of the cell. Citation: Tsang, C. C. C., P. G. J. Irwin, C. F. Wilson, F. W. Taylor, C. Lee, R. de Kok, P. Drossart, G. Piccioni, B. Bezard, and S. Calcutt (2008), Tropospheric carbon monoxide concentrations and variability on Venus from Venus Express/VIRTIS-M observations, J. Geophys. Res., 113, E00B08, doi:10.1029/2008JE003089.
1. Introduction 1.1. Past Observations of Tropospheric CO [2] The near-infrared thermal emissions were discovered by Allen and Crawford [1984] and were shown to be emanating from the deep troposphere of Venus by radiative transfer modeling by Kamp et al. [1988] and Bezard et al. [1990]. These emissions, spanning spectral wavelengths from 1.01 to 2.50 mm, are attenuated as they pass through the cloud layer, so the radiance observed from space is sensitive to changes in the total cloud optical depths. It is also sensitive to abundances of minor species, namely CO, H2O, HDO, OCS, SO2, HF, and HCl, at altitudes ranging from the surface to the base of the cloud layer at 45 km. A compre-
1 Atmospheric, Oceanic and Planetary Physics, Clarendon Laboratory, Department of Physics, University of Oxford, Oxford, UK. 2 Geophysics and Planetary Sciences, California Institute of Technology, Pasadena, California, USA. 3 LESIA, Observatoire de Paris, Meudon, France. 4 INAF-IASF, Rome, Italy.
hensive review of the near-infrared windows can be found in the work of Taylor et al. [1997] and Baines et al. [2006]. [3] While Kamp et al. [1988] were the first to propose retrieving abundances of these minor species, it was the subsequent observations by Bezard et al. [1990] and Pollack et al. [1993] using the Canada France Hawaii Telescope, who first measured the CO abundance in the troposphere, estimating mean values at 36 km of 40 ppm and 23 ± 5 ppm, respectively. Pollack et al. [1993] also retrieved a CO vertical profile, with increases of 1.20 ± 0.45 ppm km1 increasing with height. At approximately the same time, the Galileo spacecraft flew past Venus on its trajectory to Jupiter. Onboard, the NIMS imaging spectrometer took the first images from space of the nightside of Venus, most notably at the wavelengths of 1.74 and 2.30 mm. [4] Collard et al. [1993] measured the latitude distribution of CO from the Galileo/NIMS VJBAR data set. Rather than using a spectral fitting technique, Collard et al. [1993] positively detected a latitude CO gradient using a method of ratios of spectra at two different wavelengths. It was shown that a CO meridional gradient exists in the lower atmosphere, with an enhancement mainly poleward of 47°N. This work
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003089$09.00
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remained unverified until only relatively recently, when Marcq et al. [2005, 2006] using the NASA/IRTF/SpeX, observed the same CO gradient from ground-based highresolution nightside spectra. From this analysis, a lower limit for the equator-to-pole CO gradient was found to increase by approximately 15% from the equator to 40°N/S, less than the 35% increase found by Collard et al. [1993]. While this is a lower gradient, it does not cover the high latitudes covered by Galileo/NIMS. In addition, neither Collard et al. [1993] nor Marcq et al. [2006] were able to measure any zonal distribution or variability of CO because of the limited horizontal and temporal sampling of their observations. 1.2. CO as a Dynamical Tracer [5] Gierasch [1975] postulated that a mean meridional circulation would exist in the Venus atmosphere, transporting momentum from the equator to the poles. This was verified by the in situ measurements of meridional winds by the Pioneer large and small probes [Counselman et al., 1980]. Images of Venus’ dayside in the UV reveal cloud features at 65–75 km altitude; tracking these features with potential vorticity data revealed meridional wind speeds of >10 m s1 [Limaye, 2007] at this altitude. Nightside images of Venus at 1.7 and 2.3 mm exhibit contrast because of variability in the lower cloud layer at 50 km. Analysis of Galileo/NIMS [Carlson et al., 1991] found that the motion of the features indicated that meridional flow is still poleward at this altitude. This would indicate the Hadley cell extended to a depth of at least this altitude. It is this Hadley cell circulation which maintains the midlatitude jets [Lee et al., 2007], which in turn support the super-rotation, with high zonal wind speeds. In addition, because the Hadley cell is transporting angular momentum toward the poles, it may also be responsible for maintaining the polar vortex seen from a latitude of 70° poleward. There are strong indications, from both observations and general circulation models (GCM), that this equator-to-pole circulation exists in both hemispheres. However, both the vertical and latitudinal extent of the Hadley cells are yet to be fully understood. Outstanding questions include how far in latitude do the Hadley cells extend, how deep in altitude they go, are the southern and northern hemisphere cells symmetric about the equator, and are there any timedependent horizontal variations of these cells. [6] The tropospheric CO enhancement was interpreted by Taylor [1995] to be caused by the descending arm of the Hadley cell, bringing down CO from above the cloud tops where the photolysis of CO2 by UV, with an energetic threshold for photo-dissociation near 224 nm [Von Zahn et al., 1983], creates CO (noting that the CO2 cross section is dominated by Rayleigh scattering longward of 205 nm [Shemansky, 1972; Karaiskou et al., 2004]). As CO descends, it is believed to be converted to OCS and CO2 [Krasnopolsky, 2007; Fegley et al., 1997; Hong and Fegley, 1997; Y. Yung et al., Modeling the distribution of OCS in the lower atmosphere of Venus, submitted to Journal of Geophysical Research, 2008]. It is therefore possible to use CO as a dynamical tracer, which is entrained by the circulation of the lower atmosphere. This would not only provide mean abundances of CO in the lower atmosphere,
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but also provide information on the size and variability of the Hadley cell in the deep atmosphere.
2. Venus Express/VIRTIS-M Observations [7] Venus Express entered Venus orbit in April 2006. Onboard are seven science instruments devoted to observing the atmosphere and surface of Venus from UV to infrared and radio wavelengths. Visible and Infrared Thermal Imaging Spectrometer (VIRTIS), which is also flying on the Rosetta mission [Coradini et al., 1998], is one of these instruments. Covering wavelengths from 0.27 to 5.19 mm, VIRTIS is divided into two subsystems; the nonimaging H channel covering the spectral ranges from 1.84 to 4.99 mm, with a spectral resolution of 1nm, and the imaging M channel, which is further divided into two channels. One channel, viewing in the UV and visible wavelengths, is named M-Vis, while the 1.05 to 5.19 mm spectral region is sounded by the M-IR channel. The spectral resolution of VIRTIS-M is approximately 2 nm and 10 nm for the VIS and IR sub-channels, respectively. In this analysis, we will be using data from the VIRTIS-M-IR subsystem only. A comprehensive review of VIRTIS is given by Drossart et al. [2007] and Piccioni et al. [2007]. [8] While the ability of VIRTIS-H to obtain higher resolution spectra has obvious advantages, such as better retrieval of the absolute abundances of minor species as well as their vertical distribution, the multispectral imaging capabilities of VIRTIS-M makes it uniquely able to map global-scale properties such as cloud opacity and wind field measurements. It is this capability not only to measure the global-scale values of minor species, but also to do so as a function of time, which we intend to take advantage of. Indeed, this has always been one of the key measurements to be made with VIRTIS [Baines et al., 2006]. [9] The single linear array of VIRTIS-M has an instantaneous field of view of 0.25 64 mrad, divided into 256 spatial pixels. At an apogee of 66,000 km, this instantaneous field of view spans one third of the diameter of Venus. Using a scan mirror, with 256 step positions, the field of view expands to 64 64 mrad. Therefore, at apogee, a full image of the Venus disc is created when nine images are taken sequentially to create a 3 3 mosaic image. [10] Table 1 shows the list of observations used in this analysis; two orbit insertion observations, CIOB00 taken on 12 April 2006, and CIOB03 on 16 April 2006, and two nominal orbit observations in orbit number 99 and 121. The orbit insertion observations were used because the whole disc of Venus was captured in the field of view, as the initial orbit of Venus Express in April 2006 was extremely large. The distance from Venus at the time of these two observations was approximately 214,000 km and 315,000 km, respectively. Nominal orbit observations were MTP03-99-2, taken on 28 July 2006 from a distance of 61,000 km, and MTP04121-0 taken on 12 April 2006 at a distance of 65,000 km. [11] The raw observations were spectrally and spatially binned to increase the signal-to-noise ratio of the data, as well as to reduce the number of spectra needed to achieve complete spatial coverage over the observation image. As the spatial scale of each observation varied, a bin of various widths was used for different observations. A typical bin width of 20 20 pixels would be used, which is then shifted
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Table 1. Summary of the Observations Used in This Analysis and Their Important Parametersa Observation Name
Integration Time (s)
Number of Spectra
Longitude Ranges (°)
Latitude Ranges (°)
Subsolar Longitude (°)
CIOB00 – 0 CIOB03 – 7 MTP03-99-2 MTP04-121-0
8.0 8.0 8.0 3.3
300 375 300 289
0.0: 359.9 0.0: 359.9 196.71: 286.17 0.00: 360.0
0.77: 90 13.16: 89.65 22.6: 76.07 13.13: 82.69
70.74 83.69 39.71 107.69
a
To compare the local time aspect of each observation, subtracting the longitudes by the subsolar longitude will yield the local time in units of longitude.
by 10 pixels across the image for adequate Nyquist sampling. An average of some 300 spectra per observation would be generated in this manner. The associated latitudes, longitudes, and geometries is also averaged in this way. The result can be seen in an example (Figure 1) from observation CIOB00.
3. Radiative Transfer Modeling 3.1. Spectral Data [12] We employ a radiative transfer model to generate synthetic spectra, initially developed for the analysis of Galileo/NIMS data [Irwin et al., 2004], which uses the method of correlated-k [Lacis and Oinas, 1991] to calculate
the absorption coefficients of the gases in the atmosphere. The k tables are precalculated from line-by-line spectral databases. The tables are calculated on a grid of 20 evenly spaced log-pressure levels (e17 to e5 bars) and 20 evenly spaced temperature levels (150 K to 760 K). These values correspond to approximately the 120 km and surface height range. The spectral resolutions are calculated at 10 nm and 1 nm which correspond to the resolutions of the VIRTIS-M and H subsystems. For gaseous opacity, the HITEMP database is used for CO2, while the remaining trace gases use HITRAN2K. A CO2 sub-Lorenztian lineshape from Tonkov et al. [1996] was implemented, with a cutoff of 150 cm1 from the line center. CO2-CO2 self-broadening
Figure 1. A typical binning scheme used to reduce the data to manageable sizes, as well as increasing the signal-to-noise ratio. This example is taken from orbit insertion observation CIOB00. Only nightside data can be used for this analysis, hence we only use data from half of the hemisphere. 3 of 13
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Figure 2. The Venus a priori temperature profile used in the retrieval processes taken from Seiff [1983]. half-widths were taken from the HITEMP database, while the CO2-H2O foreign half-widths were taken from Delaye et al. [1989]. The other remaining gases use the air-broadened half-widths given in HITRAN2K. CO2 collision induced
absorption is parameterized according to Bezard et al. [1990], with a value of 3.5 108 cm1 amagat2 used as the opacity factor.
Figure 3. A priori concentrations of minor constituents as a function of height. 4 of 13
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3.2. Atmospheric Profiles [13] The temperature profile used as the a priori was taken from Seiff [1983] and is plotted in Figure 2. The a priori vertical distribution of gases is given in Figure 3. The amount of CO2 is 0.965 fractional abundance, which is assumed to be vertically uniform. The abundance of CO and HF has been taken from Collard [1993], where HF is uniform, and kept uniform during the retrieval process, at 0.005 ppm. The remaining vertical profiles have been taken from Marcq et al. [2005], who have taken their profiles from a mixture of Oyama et al. [1980], Von Zahn et al. [1983], and Taylor et al. [1997]. 3.3. Cloud Profile [14] We use an a priori cloud model which follows the vertical profiles of Pollack et al. [1993]. The clouds in the model extend from 50 km to some 80 km. There are four modes in the model, which are distinguished by their size parameter. The refractive indices used in the model have been taken from Palmer and Williams [1975], for a mixture of 75% sulphuric acid and 25% water vapor. This is in line with other models of the near-infrared windows such as Pollack et al. [1993] and Marcq et al. [2006]. These refractive indices are used to calculate the single scattering albedo and absorption cross sections as look-up tables assuming Mie theory. Limited tests have been conducted for different mixtures of sulphuric acid/water vapor and their impact on the 2.3 mm spectrum. Tests show, for concentrations of 75%, 85%, and 96% H2SO4, the change to the ratio value at 2.30/2.33 mm (radiance outside the CO absorption band to radiance inside the band) is on the order of much less than 1%. This shows that we are really not that sensitive to potential changes in concentrations of H2SO4/ H2O in the Venus atmosphere. [15] To account for the multiple-scattering atmosphere and aerosols, as well as to accurately model high emissionangle observations, we use a matrix operator method to calculate for multiple scattering of the radiation [Plass et al., 1973]. We calculate this using ten streams (five up, five down).
4. Analysis [16] The retrieval model used in this work, named Nonlinear Optimal Estimator for Multivariate Spectral Analysis (NEMESIS), follows the formulation of Rodgers [1976]. The retrieval scheme starts with the prescribed forward model synthetic spectrum, generated with an a priori atmosphere, described above. The convergence limit f (denoted flimit) is also required by the user to start the iterative process. This produces a predicted spectrum yn with an initial cost function fi, given by T 1 f ¼ ðym yn ÞT S1 m ðym yn Þ þ ðx xa Þ Sx ðx xa Þ
ð1Þ
where ym is the measured radiance, yn is the predicted radiance, Sm is the measurement covariance matrix, x is the state vector, Sx is the state vector covariance matrix, and xa is the a priori state vector. The measurement covariance matrix Sm contains both the measurement errors and the forward modeling errors. The forward modeling error is
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included to compensate for any lack of physics within the forward model, as well as any inaccuracies incurred when using the correlated-k approximation. The calculation and estimates of the forward modeling and a priori errors can be found in the work of Irwin et al. [2008a] and Tsang et al. [2008], respectively, and are negligible overall. [17] The first term on the right-hand side of the cost function is known as c2 and is a measure of how well the synthetic spectrum is fitting to the observed spectrum. It is f which is minimized to arrive at a possible solution. Ultimately, the optimized solution to the nonlinear problem can be found for the unknown state vector by solving xiþ1 ¼ xa þ Sx KTi Ki Sx KTi þ Sm
1
ðym yi Ki ðxa xi ÞÞ ð2Þ
where Ki is the K matrix associated with the ith iteration state vector xi. A modification is added to this GaussNewton iteration with the addition of a ‘‘braking parameter,’’ l, to keep the iteration stable for cases where the problem is nonlinear and is given by a LevenbergMarquardt constraint of the form x0iþ1 ¼ xi þ
ðxiþ1 xi Þ 1þl
ð3Þ
where x0i+1 is the modified state vector. If x0i+1 makes the fit worse, l is increased and x0i+1 is calculated again. If x0i+1 239 yields a better fit, l is decreased and the iteration continues. [18] The iterative process then proceeds to take a predefined a priori covariance error Sa, the measured spectrum ym, and the accompanied measurement covariance (error) Sm. This produces an estimate of the state vector x, with which a new spectrum yn is calculated using the radiative transfer equations. This yn has an accompanied fn. If this fn is greater than the previous f(n1), the braking parameter (equation (3)) is increased by tenfold and the iteration begins again. If fn is less than the previous f value, braking parameter l is decreased to 0.3l and the next iteration begins. This iteration scheme terminates when fn1 fn < flim it fn
ð4Þ
where flimit is a percentage change. This yields a possible state of the atmosphere xi, given the initial measured spectrum ym. In this work, the retrieval model terminates when the number of runs reaches a maximum of 60 iterations, or if the change in cost function f between iterations drops below 0.1%, whichever is reached first (these values in practice change depending on the level of fit specified and time allotted). In practice, a solution is reached usually after approximately 10– 20 iterations, when the cost function drops below the allotted 0.1% minimum. A full review of the radiative transfer and retrieval model used in this work can be found in the work of Irwin et al. [2008a].
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Figure 4. (left) The contribution functions at 2.35 mm for the background temperature, CO at nadir, and CO at an emission angle of 80°. (right) The absorbing line strengths of CO in this spectral region. [19] Figure 4, left, shows the nadir contribution functions for 2.35 mm, produced by the radiative transfer model, which shows the background emissions coming from 30 km, with a sensitivity to CO at 35 km. Between 2.30 and 2.5 mm, the rotational-vibrational lines of the 2-0 band of carbon monoxide are active [Bezard et al., 1990], the spectral lines of which are also shown in Figure 4, right. It is these lines which provide us with the ability to retrieve the concentration of CO in the 35 km height range. [20] We take the entire spectral range, from 2.20 to 2.50 mm, to retrieve the tropospheric carbon monoxide concentrations. A sample spectrum can be seen in Figure 5. From wavelengths of 2.18 to 2.30 mm, the absorption is dominated by the presence of CO2. The radiation is also attenuated by the intervening cloud layer. This attenuation has a predominantly gray effect over this entire spectral window [Bezard et al., 1990; Marcq et al., 2006; Pollack et al., 1993]. Thus, the spectra covering these wavelengths are used to fit to the varying optical depths of the clouds across the planet. Once this is done, the rest of the spectra, from 2.30 to 2.50 mm, are assumed to have the same absorption as that caused by the clouds from 2.18 to 2.30 mm. [21] The remaining absorption in this region is due to the presence of trace gases which are active absorbers in this spectral region. The sensitivity of this spectrum in wavelength and height of trace gases can be found in the work of
Tsang et al. [2008] and Baines et al. [2006]. While we fit the 2-0 band of CO from 2.30 to 2.35 mm, we also simultaneously fit the rest of the spectrum to retrieve abundances for H2O and OCS, the results of which will not be shown in this work as it has no impact on the retrieved CO abundances. An example of our ability to fit our synthetic spectra to the data is shown in Figure 5. The residual, a simple difference between the measured and the synthetic spectrum, is also shown in Figure 5. Once f has been minimized, a best estimate for the CO abundance is found. In the case of CO, we retrieve a scaling factor which is multiplied by the a priori profile to yield a new abundance. Since the peak of the contribution function for CO is at approximately 35 km, we will show the abundances for that altitude.
5. Results [22] Figure 6 shows the observation and the results from the first orbit insertion observation, CIOB00. The 2.3 mm radiances resulting from the variations in the optical depths of the clouds that are attenuating the thermal radiation from above 35 km are shown at the spatial resolution of the binned spectra. After spectral fitting using the method described above, we can also display the c2/n values for different areas of the map, showing our ability to fit spectra
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Figure 5. An example Visible and Infrared Thermal Imaging Spectrometer (VIRTIS)-M spectrum between 2.18 and 2.50 mm used in this analysis. (top) Solid line (no symbol) is the measured spectrum, with dotted lines showing the limits of uncertainty in the radiances which arise from forward modeling and the measurement of the radiances. Solid, star symbol line is the model fit when fitting to cloud opacity. Solid, diamond symbol line is the model fit when fitting to CO, H2O, and OCS, assuming a cloud opacity derived from the solid star line. (bottom) The radiance residual between the measured and fitted spectrum. at those precise locations. The c2/n map is quite homogeneous, with values of between 1 and 3. Much higher c2/n values appear near the day-night terminators and at the poles because of bad spectra and poor signal-to-noise. It should be noted here that the c2/n fitting is still in an acceptable range, even though the optical depth has increased because of the presence of the polar cloud and vortex. The primary driver of the retrieved errors is the cloud opacity. If the cloud opacity is extremely high, like at the polar vortex, the signal-to-noise reduces dramatically, and thus the retrieved errors increase as well. [23] The resulting CO abundances and related errors are plotted on Figure 6. The mean retrieved errors in the abundances are of the order of 2 to 3 ppm (approximately 10% of the retrieved variance to mean ratio abundances) in the midlatitudes. The errors then increase toward the poles and the equator. The large errors at the equator are due to the geometry of the observation during this orbit insertion phase, where the subspacecraft point was above the poles. This results in radiances with high emission angles at the equator, with reduced signal-to-noise ratios. Correspondingly, the polar regions beyond 70°S are also subject to low signal-to-noise because of the increase in optical depths of the polar clouds. [24] We see that the mean equatorial values of 23 ± 3 ppm increase toward the poles, reaching a maximum near 60°S, of 32 ± 3 ppm. The distribution of CO poleward of 60°S
also indicates a possible decrease in abundance. This general first-order trend of increasing from equator, reaching maximum near/at 60°S, and inversion toward the poles is best seen in Figure 6. This map also shows a degree of variability of CO in solar longitude. We can best see this in Figure 7, where we plot the retrieved CO abundances as a function of the solar longitude. Solar longitude of 100° corresponds to the dusk terminator, while 260° corresponds to the dawn terminator. The first-order trends are apparent, but variability seen in Figure 7 as a function of solar longitude indicates a number of general features which are of note: (1) The mean abundance of CO on longitudes near the dusk terminator has greater absolute value than post local midnight and near dawn. (2) The maximum of the CO abundance near dusk is at 60°S, while the CO maximum near dawn is closer to 70°S. (3) Midlatitude (approximately 40°S) abundances of CO at approximately 26 ± 2 ppm have less variation than at other latitudes. [ 25 ] The same picture is painted with observation CIOB03, taken some 96 h later, again from orbit insertion (Figure 8). The same evening/morning asymmetry in the CO abundance can be seen, although it is less pronounced in this observation than CIOB00. Less of the equator is seen as CIOB03 was taken in the ascending branch of the Venus Express orbit. We also note that the midlatitude values are very constant across solar longitude, while variability can be seen in the dusk and dawn part of the nightside disc. It
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Figure 6. Orbit Insertion Observation CIOB00: (top left) 2.30 1 m radiance, a proxy for cloud opacity, plotted after binning; (top right) c2/n values when fitting for CO between 2.30 and 2.35 mm; (bottom left) the retrieved CO abundance at 35 km; and (bottom right) the error in the retrieved CO abundance. Areas of the observation where the retrieved error exceeds 8 ppm are colored white. Regions of high c2/n are reflected in the map of the retrieved errors (in white).
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Figure 7. The retrieved CO gradient versus latitude from CIOB00 observation, plotted as a function of solar longitude, interpolated in spacings of 5°. These retrieved CO abundances are referenced to 35 km. Dusk terminator at 100° and dawn terminator at 260°. should also be noted that the mean abundances at 60°S for CIOB00 and CIOB03 are 32 and 34 ± 2 ppm, respectively. [26] To investigate some of these features further, to enhance the sampling and to see the if the degree of variability of CO concentrations changes over time, we take observations from the nominal orbit during ‘‘case 2’’ observations, MTP03-99-2 and MTP04-121-0. These are shown in Figures 9 and 10, respectively. MTP03-99-2 shows a constant band of maximum CO abundance of between 35 to 36 ± 2 ppm at a latitude of 60°S. Zonal variability at this latitude can just be made out, with local maxima of CO at 90 and 135°W and a minimum at approximately 112°W, with a value of 34 ppm. However, it should be noted that the retrieved error is of the order of these fluctuations and therefore is not entirely without doubt. In addition, the maximum in CO commencing at 60°S continues at a steady value until as high as 70°, where a slight decrease in the abundance past 70°S can be made out. The retrieved errors in this region have not increased significantly. Observation MTP04-121-0 shows a similar trend. Again, the CO maximum is located at 60°S, with mean upper limit of 36 ± 3 ppm, but in this instance seems more localized around a longitude of 45°W, decreasing by 3 ppm toward 68°W. However, caution should be taken again as this decrease is
of the order of the retrieved error. The decrease from 65°S in the CO abundance to 30 ppm toward the poles is matched by increased error caused by the increase in the cloud optical depth of the polar vortex, decreasing the signal-tonoise ratio. Therefore, we cannot fully confirm the validity of the overturning from this image alone.
6. Discussion [27] We first confirm the equator-to-pole increase of CO in the troposphere first seen by Collard et al. [1993] from Galileo/NIMS data, validated by Marcq et al. [2006] from IRTF/SpeX data. It is also consistent with observations made by Marcq et al. [2008], using much higher spectral resolution subsystem of VIRTIS-H, who have shown very similar results for both equatorial and polar values of CO. The authors of the above work found a 30 ± 15% in- crease from 0° to 60°, corresponding to 24 ± 3 to 31 ± 2 ppm at 36 km. This is very consistent with the results of this paper, which show on first order, an increase in the same latitude range of 23 ± 2 to 32 ± 2 ppm at the same altitude. [28] In addition, because of the unique imaging capabilities of the VIRTIS-M subsystem, this analysis has also possibly revealed two new features in the tropospheric CO
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Figure 8. Orbit Insertion Observation CIOB03: (left) Retrieved CO abundance at 35 km with (right) the associated retrieval errors. Areas of the observation where the retrieved error exceeds 8 ppm are colored white.
Figure 9. (left) The retrieved CO abundance at 35 km and (right) the errors for observation MTP03-99-2. Variability in the CO abundance at 60°S across different longitudes can be seen. 10 of 13
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Figure 10. (left) Retrieved CO from observation MTP04-121-0, with (right) the retrieved errors. Areas of the observation where the retrieved error exceeds 8 ppm are colored white. distribution which have not been previously observed and warrant further discussion: the evening/morning asymmetry in the CO zonal profiles and the reversal in CO concentration poleward of 60°. Both these features can be explained by a combination of the Hadley-type circulation, as depicted in Figure 11, and photochemistry from the cloud tops.
[29] In regard to the overturning of CO past 60°S, the mechanism for this is likely to be purely dynamical in nature. The general idea is that the descending branch of the Hadley cell occurs at 60°. This is consistent with GCM models such as Lee et al. [2007] where the maximum downwelling occurs at 60°N/S. It is also consistent with measurements of the dayside meridional winds from Pioneer
Figure 11. A pictorial representation of the Hadley cell type circulation which exists in the Venus atmosphere. A complex interplay between dynamical and chemical cycles combines to yield the features seen in the retrieved CO maps. Photo-disassociation reaction taken from Huebner et al. [1992]. 11 of 13
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Venus which indicate the horizontal convergence of the meridional flow is greatest between 50 and 60° [Limaye, 2007]. We expect the maximum downwelling of the Hadley cell to be associated with regions of maximum vertical winds, and therefore would bring the most CO rich air from mesosphere (>65 km) to the troposphere. If this downwelling is greatest at 60°, it implies a poleward decrease in CO concentration past 60°S, and therefore would be consistent with our observations. [30] The situation with respect to the east-west asymmetry in the CO abundances is probably a mixture of photochemistry and dynamics processes, which is responsible for the concentration of CO at this altitude since the dominant chemistry occurs way above or way below this 35 km region. The supply of CO is governed by the upper atmosphere source. The two options to create this east-west asymmetry in CO at 35 km is either an east-west gradient in the source of CO in the source region and constant downwelling, or a constant source of CO in the source region and an east-west gradient in the downwelling across longitudes. It is indeed likely to be a combination of both processes. [31] The evening/morning asymmetry seen in Figures 6 and 8 can be, on first order, simply a function of the photochemistry which is occurring at the cloud tops at 65 km. If we consider what is occurring at the cloud tops, during the early morning, we would expect a minimum in the early morning with a ramp up of CO production in the morning sunshine due to the photolysis by solar UV, and would depend on the photolysis rate of CO above the clouds. This would peak in the early evening (before sunset probably). Maps of CO on the nightside of Venus above the clouds [Irwin et al., 2008b] indicate enhanced CO at the evening terminator and are consistent with these observations of the tropospheric CO and photochemical mechanisms above the clouds. As the photolysis stops because of the lack of sunlight, no more CO is produced on the nightside. The meridional circulation, however, will continue to pump mesospheric air into the lower troposphere. At the start of the evening you would therefore expect high abundances of CO in the lower atmosphere at the location of the downwelling of the Hadley cell and for this concentration to decay as the night progresses, assuming the vertical velocity of this downwelling was constant across the nightside (which is most likely not the case). However, CO is constantly being destroyed by sink reactions, most notably by OCS [Pollack et al., 1993] and S2 [Hong and Fegley, 1997], and reconverted back to CO2. Therefore, we also need to know something about the creation and destruction rates for CO in the lower atmosphere, as well as the meridional overturning rates (10 m s1 poleward winds at 100 mbar). From the observations of Figures 6 and 8, a change in the concentration of CO of 2 ppm from the evening to morning terminator can be seen. There may be some problem accounting for this given what is known about the reaction rates and atmospheric dynamics.
7. Conclusions [32] Observations of the nightside of Venus taken from Venus Express/VIRTIS-M instrument at wavelengths between 2.20 to 2.50 mm have been used to measure and map the distribution of CO in the troposphere of Venus. We have
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taken spectra from four different observations, two from orbit insertion, and two from the nominal orbit, which are then coadded and binned for increased signal-to-noise and homogeneous spatial sampling. We use a spectral fitting algorithm to first account for variations in the cloud optical depth by fitting spectra from 2.20 to 2.30 mm. This is then followed by fitting the remaining spectra from 2.30 to 2.35 mm, which is sensitive to the 2-0 CO band, by scaling a known a priori profile. [33] We first confirm the presence of the CO enhancement from equator to pole, with a mean increase of between 40% and 45% from 23 ± 2 ppm to 32 ± 2 ppm, which is consistent with observations made by the Galileo/NIMS flyby of Venus in 1990. We find the peak value of CO to occur between 55 and 65°S. We also find evidence of two possibly new features in this CO gradient; decreases of CO poleward of 60°S and variability in longitude. [34] A mixture of photochemistry and dynamics will be responsible for the concentration of CO at this altitude because the dominant chemistry does not occur near 35 km. Advection of CO from below is unlikely because of a lack of a continuous source, leaving photolysis of CO2 above 60 km as the most likely source of CO in the lower atmosphere. The CO is then advected into the lower atmosphere, particularly in the downwelling region near 60°, where it is destroyed. Given the need for sunlight to photolyse the CO2, the distribution of CO at the source altitude probably follows the solar zenith angle, peaking in the afternoon hemisphere because of the background flow. Poleward advection in the source region and downwelling at the poles then transports the CO into the lower atmosphere, producing the observed latitudinal gradient at 35 km. These observations would then place the downwelling branch of the Hadley cell near 60°, consistent with observations of peak westward flow at 60° and GCM simulations such as Lee et al. [2007]. It is also consistent with measurements of the dayside meridional winds from Pioneer Venus which indicate the horizontal divergence (converging air mass) of the meridional flow is greatest between 50 and 60° [Limaye, 2007]. Poleward of 60°, the vortex ‘‘barrier’’ and decreasing sunlight would contribute to a negative latitudinal gradient in CO, as observed. In addition, the fact that we observe the CO gradient at 35 km, which is a proxy for the return branch of the Hadley cell, implies the vertical extent of the cell(s) is at least 40 km. [35] Acknowledgments. I would like to thank Emmanuel Marcq for his discussions of CO and his work with VIRTIS-H spectra and its findings. We would also like to thank Frank Mills and Robert Carlson for their discussions on this subject. This work was made possible from funding given by the United Kingdom Science Technology Facilities Council, as well as the continued support of CNES and ASI.
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spectroscopy measurements at Venus, Science, 253, 1541 – 1548, doi:10.1126/science.253.5027.1541. Collard, A. (1993), Thermal emission from the nightside of Venus, D.Phil. thesis, Univ. of Oxford, UK. Collard, A., F. Taylor, S. Calcutt, R. Carlson, L. Kamp, and K. Baines (1993), Latitudinal distribution of carbon monoxide in the deep atmosphere of Venus, Planet. Space Sci., 41(7), 487 – 494. Coradini, A., et al. (1998), VIRTIS: An imaging spectrometer for the Rosetta mission, Planet. Space Sci., 46(9 – 10), 1291 – 1304. Counselman, C., S. Gourevitch, R. King, G. Loriot, and E. Ginsberg (1980), Zonal and meridional circulation of the lower atmosphere of Venus determined by radio interferometry, J. Geophys. Res., 85, 8026 – 8030, doi:10.1029/JA085iA13p08026. Delaye, C., J. Hartmann, and J. Taine (1989), Calculated tabulations of H2O line broadening by H2O, N2, O2, and CO2 at high temperature, Appl. Opt., 28(23), 5080 – 5087. Drossart, P., et al. (2007), Scientific goals for the observation of Venus by VIRTIS on ESA/Venus Express mission, Planet. Space Sci., 55, 1653 – 1672, doi:10.1016/j.pss.2007.01.003. Fegley, B. J., M. Y. Zolotov, and K. Lodders (1997), The oxidation state of the lower atmosphere and surface of Venus, Planet. Space Sci., 125, 416 – 439. Gierasch, P. (1975), Meridional circulation and the maintenance of the Venus atmospheric rotation, J. Atmos. Sci., 32, 1038 – 1044, doi:10.1175/ 1520-0469(1975)032<1038:MCATMO>2.0.CO;2. Hong, Y., and B. Fegley (1997), Formation of carbonyl sulphide (OCS) from carbon monoxide and sulphur vapour and applications to Venus, Icarus, 130, 495 – 504, doi:10.1006/icar.1997.5824. Huebner, W. F., J. J. Keady, and S. P. Lyon (1992), Solar photon rates for planetary atmospheres and atmospheric pollutants, Astrophys. Space Sci., 195, 1 – 294, doi:10.1007/BF00644558. Irwin, P., P. Parrish, T. Fouchet, S. Calcutt, F. Taylor, A. Simon-Miller, and A. Nixon (2004), Retrievals of Jovian tropospheric phosphine from Cassini/CIRS, Icarus, 172, 37 – 49, doi:10.1016/j.icarus.2003.09.027. Irwin, P. G. J., N. Teanby, R. de Kok, L. Fletcher, C. Howett, C. C. C. Tsang, C. F. Wilson, S. B. Calcutt, C. Nixon, and P. Parrish (2008a), The NEMESIS planetary atmosphere radiative transfer and retrieval tool, J. Quant. Spectrosc. Radiat. Transf., 109, 1136 – 1150, doi:10.1016/ j.jqsrt.2007.11.006. Irwin, P. G. J., R. de Kok, A. Negra˜o, C. C. C. Tsang, C. F. Wilson, P. Drossart, G. Piccioni, D. Grassi, and F. W. Taylor (2008b), Spatial variability of carbon monoxide in Venus’ mesosphere from Venus Express/VIRTIS Visible and Infrared Thermal Imaging Spectrometer measurements, J. Geophys. Res., 113, E00B01, doi:10.1029/ 2008JE003093. Kamp, L., F. Taylor, and S. Calcutt (1988), Structure of Venus’s atmosphere from modelling of night-side infrared spectra, Nature, 336, 360 – 362, doi:10.1038/336360a0. Karaiskou, A. C., C. Vallance, V. Papadakis, I. M. Vardavas, and T. P. Rakitzis (2004), Absolute absorption cross-section measurements of CO2 in the ultraviolet from 200 to 206 nm at 295 and 373 K, Chem. Phys. Lett., 400, 30 – 34. Krasnopolsky, V. A. (2007), Chemical kinetic model for the lower atmosphere of Venus, Icarus, 191, 25 – 37, doi:10.1016/j.icarus.2007.04.028. Lacis, A., and V. Oinas (1991), A description of the correlated k-distribution method for modeling non-grey gaseous absorption, thermal emission and multiple scattering in vertically inhomogeneous atmospheres, J. Geophys. Res., 96, 9027 – 9063, doi:10.1029/90JD01945. Lee, C., S. Lewis, and P. Read (2007), Superrotation in a Venus general circulation model, J. Geophys. Res., 112, E04S11, doi:10.1029/ 2006JE002874. Limaye, S. (2007), Venus atmospheric circulation: Known and unknown, J. Geophys. Res., 112, E04S09, doi:10.1029/2006JE002814.
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Marcq, E., B. Bezard, T. Encrenaz, and M. Birlan (2005), Latitudinal variations of CO and OCS in the lower atmosphere of Venus from nearinfrared nightside spectro-imaging, Icarus, 179, 375 – 386, doi:10.1016/ j.icarus.2005.06.018. Marcq, E., T. Encrenaz, B. Bezard, and M. Birlan (2006), Remote sensing of Venus’ lower atmosphere from ground-based spectroscopy: Latitudinal and vertical distribution of minor species, Planet. Space Sci., 54, 1360 – 1370, doi:10.1016/j.pss.2006.04.024. Marcq, E., B. Bezard, P. Drossart, and G. Piccioni (2008), A latitudinal survey of CO, OCS, H2O and SO2 in the lower atmosphere of Venus: Spectroscopic studies using VIRTIS-H, J. Geophys. Res., doi:10.1029/ 2008JE003074, in press. Oyama, V., G. Carle, F. Woeller, J. Pollack, R. Reynolds, and R. Craig (1980), Pioneer Venus gas chromatography of the lower atmosphere of Venus, J. Geophys. Res., 85, 7891 – 7902, doi:10.1029/JA085iA13p07891. Palmer, K., and D. Williams (1975), Optical constants of sulfuric acid; Application to the clouds of Venus?, Appl. Opt., 14(1), 208 – 219. Piccioni, G., et al. (2007), VIRTIS: The Visible and Infrared Thermal Imaging Spectrometer, ESA Spec. Publ. SP-1295, pp. 1 – 27, Eur. Space Agency, Paris. Plass, G. N., G. W. Kattawar, and F. E. Catchings (1973), Matrix operator theory of radiative transfer: 1. Rayleigh scattering, Appl. Opt., 12, 314 – 329. Pollack, J., J. Dalton, D. Grinspoon, R. Wattson, R. Freedman, D. Crisp, D. Allen, B. Bezard, C. de Bergh, and L. Giver (1993), Near-infrared light from Venus’ nightside: A spectroscopic analysis, Icarus, 103, 1 – 42, doi:10.1006/icar.1993.1055. Rodgers, C. (1976), Retrieval of atmospheric temperature and composition from remote measurements of thermal radiation, Rev. Geophys., 14(4), 609 – 624, doi:10.1029/RG014i004p00609. Seiff, A. (1983), Thermal structure of the atmosphere of Venus, in Venus, pp. 215 – 279, Univ. of Ariz., Tucson. ˚, Shemansky, D. E. (1972), CO2 extinction coefficient 1700 – 3000 A J. Chem. Phys., 56, 1582 – 1587, doi:10.1063/1.1677408. Taylor, F. (1995), Carbon monoxide in the deep atmosphere of Venus, Adv. Space Res., 16(6), 81 – 88, doi:10.1016/0273-1177(95)00253-B. Taylor, F., D. Crisp, and B. Bezard (1997), Near-infrared sounding of the lower atmosphere of Venus, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, pp. 325 – 351, Univ. of Ariz. Press, Tucson. Tonkov, M., N. Filippov, V. Bertsev, J. Bouanich, V.-T. Nguyen, C. Brodbeck, J. Hartmann, C. Boulet, and F. Thibault (1996), Measurements and empirical modeling of pure CO2 absorption in the 2.3 mm region at room temperature: Far wings, allowed and collision- induced bands, Appl. Opt., 35(24), 4863 – 4870. Tsang, C. C. C., P. G. J. Irwin, F. W. Taylor, and C. F. Wilson (2008), A correlated-k model of radiative transfer in the near infrared windows of Venus, J. Quant. Spectrosc. Radiat. Transf., 109, 1118 – 1135, doi:10.1016/ j.jqsrt.2007.12.008. Von Zahn, U., S. Kumar, H. Niemann, and R. Prinn (1983), Composition of the Venus atmosphere, in Venus, pp. 299 – 430, Univ. of Ariz., Tucson.
B. Bezard and P. Drossart, LESIA, Observatoire de Paris, 5 place Jules Janssen, F-92195 Meudon, France. S. Calcutt, R. de Kok, P. G. J. Irwin, F. W. Taylor, C. C. C. Tsang, and C. F. Wilson, Atmospheric, Oceanic and Planetary Physics, Clarendon Laboratory, Department of Physics, University of Oxford, Parks Road, Oxford OX1 3BH, UK. (
[email protected]) C. Lee, Geophysics and Planetary Sciences, California Institute of Technology, MC 150-21, Pasadena, CA 91125, USA. G. Piccioni, INAF-IASF, Via del Fosso del Cavaliere 100, I-00133 Rome, Italy.
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Spatial variability of carbon monoxide in Venus’ mesosphere from Venus Express/Visible and Infrared Thermal Imaging Spectrometer measurements P. G. J. Irwin,1 R. de Kok,1 A. Negra˜o,2 C. C. C. Tsang,1 C. F. Wilson,1 P. Drossart,3 G. Piccioni,4 D. Grassi,2 and F. W. Taylor1 Received 1 February 2008; revised 16 April 2008; accepted 5 May 2008; published 23 July 2008.
[1] Observations of Venus’ mesosphere by the Visible and Infrared Thermal Imaging
Spectrometer (VIRTIS)-M instrument of Venus Express have been used to investigate the spatial distribution of CO above Venus’ nightside cloud tops by fitting the CO absorption in the (1–0) CO band around 4.7 mm. We find little spatial variation in the abundance of CO at midlatitudes, with a retrieved abundance of approximately 40 ± 10 ppm just above the cloud tops between 65 and 70 km altitude. Unfortunately, we find it very difficult to constrain the abundance of CO in the cold polar collar, centered at about 70°S, as the retrieved temperature structure in the CO line-forming region masks the absorption lines. However, there is a possibility that CO increases toward the poles, as we detect a significant signature of high levels of CO over Venus’ south polar dipole feature in all the observations analyzed so far. To constrain the abundance of CO more closely will require the analysis of higher-resolution VIRTIS-H observations. In addition, limb observations would greatly help to resolve any possible temperature/cloud ambiguities and allow us to assess vertical variations in the abundance of CO. Citation: Irwin, P. G. J., R. de Kok, A. Negra˜o, C. C. C. Tsang, C. F. Wilson, P. Drossart, G. Piccioni, D. Grassi, and F. W. Taylor (2008), Spatial variability of carbon monoxide in Venus’ mesosphere from Venus Express/Visible and Infrared Thermal Imaging Spectrometer measurements, J. Geophys. Res., 113, E00B01, doi:10.1029/2008JE003093.
1. Introduction [2] Carbon monoxide is produced at equatorial and midlatitudes in the dayside upper atmosphere of Venus by UV photolysis (l < 224 nm) of CO2 [von Zahn et al., 1983; Huebner et al., 1992]. It is thought to be transported toward the poles by the planet’s Hadley cell circulation, where it is then subducted and recombined with O near Venus’ hot surface, perhaps via sink reactions with either carbonyl sulfide (OCS) [Pollack et al., 1993], or S2 [Hong and Fegley, 1997]. It is thus expected that the volume mixing ratio of CO decreases as we move deeper into the atmosphere, both owing to such recombination reactions and also owing to the CO-rich air from high altitudes being steadily diluted by mixing with CO-poor air at lower altitudes. Superimposed on this meridional circulation, CO is also rapidly transported longitudinally in Venus’ upper atmosphere by Venus’ strong zonal winds at these pressure levels. Thus, even though it is produced on Venus’ dayside, CO is expected to be quickly transported to Venus’ night-
1 Atmospheric, Oceanic, and Planetary Physics, Clarendon Laboratory, University of Oxford, Oxford, UK. 2 INAF-IFSI, Rome, Italy. 3 LESIA, Observatoire de Paris, Meudon, France. 4 INAF-IASF, Rome, Italy.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003093$09.00
side where it can be detected at thermal-IR wavelengths in the (1 –0) CO band around 4.7 mm. [3] Existing observations of CO prior to the arrival of the Venus Express mission at Venus in 2006 are reviewed by de Bergh et al. [2006] and issues concerning the modeling of the distribution of CO are reviewed by Mills and Allen [2007]. The abundance of CO at high altitudes (around 100 km altitude) is estimated from observations at millimeter [e.g., Clancy and Muhleman, 1991; Gurwell et al., 1995] and submillimeter [e.g., Clancy et al., 2003] wavelengths to be highly variable and of the order of 100 – 1000 ppm, while deeper in the atmosphere, the CO abundance has been inferred from measurements of the near-infrared emission from Venus’ night side. Be´zard et al. [1990] and Pollack et al. [1993] used data from the Canada France Hawaii Telescope to infer mean values of 40 ppm and 23 ± 5 ppm at 36 km respectively, with Pollack et al. [1993] also estimating that the abundance of CO increased with height at a rate of 1.20 ± 0.45 ppm/km. These observations roughly coincided with the flyby of Venus of NASA’s Galileo spacecraft, on its way to Jupiter. During the encounter, the NIMS (Near-Infrared Mapping Spectrometer) instrument made the first space-based observations of Venus’ night side near-infrared emission and Collard et al. [1993] used these data to determine the latitudinal distribution of CO in the deep atmosphere, finding that it increased toward the poles with a maximum abundance at approximately 60°N.
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Table 1. Summary of Observations Used in This Studya Name
Orbit
VI0029_06 VI0029_07 VI0029_08 VI0038_00 VI0067_00 VI0067_01 VI0067_02 VI0067_03 VI0067_04 VI0067_05
29 29 29 38 67 67 67 67 67 67
Date 19 19 19 28 26 26 26 26 26 26
May May May May June June June June June June
2006 2006 2006 2006 2006 2006 2006 2006 2006 2006
Start Time, UT
End Time, UT
nl
ns
nl
texp (s)
1658 1713 1729 1641 1449 1503 1517 1534 1542 1553
1708 1724 1740 1651 1454 1509 1523 1540 1548 1559
432 432 432 432 432 432 432 432 432 432
256 256 256 256 256 256 256 256 256 256
253 254 254 228 126 126 126 126 126 126
0.36 0.36 0.36 0.36 0.36 0.36 0.36 0.36 0.36 0.36
a
Here nl is the number of wavelengths, ns is the number of spatial samples, and nl is the number of spatial lines; texp is the exposure time.
[4] New observations of Venus’ near-infrared emission have been made by the VIRTIS instrument on Venus Express and these data have been used to investigate further the spatial variation of CO in the deep atmosphere by C. C. C. Tsang et al. (Tropospheric Carbon Monoxide Concentrations and Variability on Venus from Venus Express/VIRTIS-M Observations, submitted to Journal of Geophysical Research, 2008) and E. Marcq et al. (A latitudinal survey of CO, OCS, H2O and SO2 in the lower atmosphere of Venus: Spectroscopic studies using VIRTIS-H, submitted to Journal of Geophysical Research, 2008). Higher in the atmosphere, SPICAV/SOIR UV data have been used to determine the vertical distribution of CO above 60 km, finding values of 10 –20 ppm at midlatitudes at an altitude of 70 km, decreasing to 4 ppm at 90 km, and then increasing rapidly at higher altitudes [Svedhem et al., 2007; Bertaux et al., 2007].
2. Measurements [5] The Venus Express spacecraft was launched on a Soyuz-Fregat rocket from Baikonur, Kazakhstan on 9 November 2005 and went into a near-polar 24 h orbit about Venus on 11 April 2006, where it will remain until at least May 2009. [6] The spacecraft carries a number of remote sensing instruments including VIRTIS (Visible and Infrared Thermal Imaging Spectrometer), which covers the wavelength range 0.27 to 5.19 mm. The VIRTIS instrument is split into two main subsystems, the high-resolution subsystem, VIRTIS-H, and the mapper subsystem, VIRTIS-M. The VIRTIS-H subsystem is an Echelle spectrometer covering the 1.84 to 4.99 mm range at a spectral resolution of 0.001 mm, with an instantaneous field of view (IFOV) of 0.45 2.25 mrad. The VIRTIS-M subsystem, has an IFOV of 0.25 64 mrad sampled by 256 rows of a CCD array, which can be scanned to generate an image covering 64 64 mrad with an angular resolution of 0.25 0.25 mrad. At an apogee of 66,000 km, Venus’ disk extends approximately 180 mrad and thus a complete image of Venus’ disk may be generated by mosaicking 3 3 VIRTIS-M observations. The VIRTIS-M subsystem is itself split into two components: one component covers the 0.27 to 1.0 mm wavelength range with a spectral resolution of 0.002 mm, while an infrared component covers the 1.05 to 5.19 mm range at a lower spectral resolution of 0.01 mm. [7] To investigate the distribution of CO in Venus’ mesosphere we analyzed VIRTIS spectra from 4– 5 mm.
In this spectral range, there is a strong CO2 absorption band at 4.3 mm and a weaker one at 4.8 mm, with the (1 – 0) absorption band of CO observable at around 4.7 mm. No other gases have a significant contribution in this range. At other wavelengths where gas absorption becomes negligible, the absorption of the upper cloud decks becomes important. Day side observations at these wavelengths are considerably complicated by reflected sunlight from the cloud tops and so we chose to concentrate upon VIRTIS-M night side observations where only thermal emission is important, and for which there exist several wide area maps from the south pole to the equator. We chose two sets of observations in particular, VI0029 and VI0067, whose characteristics are summarized in Table 1. These mapping observations were chosen because, at the time of writing, they covered the largest area and had the longest integration time, and consequently best signalto-noise ratio. The signal-to-noise ratio of these spectra typically varies from about 1.0 in the center of the strong CO2 band at 4.3 mm to roughly 40 where the radiance is greatest, near 4.9 mm. To improve the signal-to-noise ratio, the 4.0– 4.95 mm spectra from the data ‘cubes’ sampling Venus’ night side for these orbits (six cubes for VI0067 and three for VI0029) were averaged into 20 20 pixel boxes and stepped by 10 pixels in each direction to achieve Nyquist sampling. The areas covered by both sets of observations are shown in Figure 1. pThe ffiffiffi error on the spectral points was set to either the NESR/ n or the variance of the measured points pffiffiffi divided by n, whichever was larger, where n is the number of points averaged in a single bin and NESR is the noise equivalent spectral radiance. [8] We also searched the measured data set for VIRTIS-M limb-sounding observations of Venus’ night side, but at the time of writing could only identify one suitable observation, which was of the equatorial region and had insufficient vertical sampling to improve upon what could be extracted from the VIRTIS-M nadir observations.
3. Retrieval Model [9] To analyze the recorded spectra, we used the NEMESIS correlated-k radiative transfer and retrieval model [Irwin et al., 2008], which has been successfully applied to numerous planetary remote sensing missions, and has also been used by Tsang et al. (submitted manuscript, 2008) to determine the variability of CO in Venus’ deep atmosphere. The NEMESIS retrieval code is based upon the optimal estimation formalism of Rodgers [2000], which assumes that a good knowledge exists of the expected
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Figure 1. Sampling of Venus’ southern hemisphere by the Venus Express/VIRTIS (left) VI0067 observations and (right) VI0029 observations. variation in the retrieved parameters from other sources. Such a situation does not exist in this case and hence the a priori guessed profiles and covariance matrices are instead tuned to provide a reasonable balance between precision and vertical smoothing, in the manner of the constrained linear inversion technique [Conrath et al., 1998; Hanel et al., 2003], as described by Irwin et al. [2008]. [10] To use the model, k-distribution tables were first generated with VIRTIS-M resolution from line data for the different spectrally active gases in this spectral and pressure region. We used line data from the HITRAN 2004 database [Rothman et al., 2005] with a Voigt line shape for all gases except CO2, for which we used a sub-Lorentzian line shape of Tonkov et al. [1996] with a cutoff at 150 cm1 from the line centers. Although this line shape was originally developed for analyzing spectra at 2.3 mm, we find that it allows us to reproduce the entire 4 – 5 mm spectrum very well, in particular the section from 4 to 4.3 mm, which is not wellproduced by other line shapes. Our final retrieved temperature maps match closely those retrieved by Grassi et al. [2008] from the 4.3 to 5.1 mm range (excluding the CO band at 4.7 mm) who use a different sub-Lorentzian correction. Hence, the use of the Tonkov et al. [1996] line shape does not significantly effect our temperature retrievals, and allows us to better model the weak CO absorption feature. The k-tables were calculated on a grid with 20 temperatures spread equally between 100 and 350 K, and 20 pressure levels spread logarithmically between 4 108 and 2.7 bar, with a square spectral resolution of width 10 nm. When calculating spectra, NEMESIS convolved the output with a further square bin of width 10 nm, to achieve an effective instrument function close to that experimentally determined for VIRTIS-M. [11] For these tests we assumed that the upper cloud optical depth was dominated by a lognormal distribution of H2SO4 droplets of mean radius of 1.0 mm and a variance of 1.29 (i.e., the mode 2 particles of Grinspoon et al. [1993]) and computed the extinction cross-section spectrum from Mie theory. Scattering effects were found to be negligible and so synthetic spectra were calculated assuming thermal emission only. [12] Our first-guess, or a priori, temperature profile was taken from Seiff [1983], while for clouds we assumed the
model of Roos et al. [1993], who used limb-darkening observations from Galileo/NIMS observations to determine that the mean cloud structure at equatorial latitudes was best matched by a cloud with a scale height of approximately 5.2 km. The a priori cloud profile density was thus set to decrease with height at this rate and the profile scaled to give a visible optical depth, at 630 nm, of 1.0 at an altitude of 65 km. For gaseous profiles the mole fraction of CO2 was set to 0.965 [von Zahn et al., 1983] and the remaining gases included were H2O and SO2, whose a priori profiles were set as recommended by von Zahn et al. [1983] and Krasnopolskii and Parshev [1983], although these gases were found to contribute negligibly to this spectral region. For CO we assumed a constant mole fraction with an a priori abundance of 100 ppm, based on determinations of the abundance in the lower atmosphere of 30 ppm (Tsang et al., submitted manuscript, 2008) and assuming the abundance increases with height. [13] NEMESIS can simultaneously retrieve multiple variables and can model atmospheric constituents either as continuous or parameterized profiles. To assess the retrieval model and the validity of the correlated-k approximation [Lacis and Oinas, 1991], NEMESIS was first used to retrieve the vertical temperature profile alone from VIRTIS-M observations and the results compared with the temperature retrievals of Grassi et al. [2008]. Very good agreement was achieved, adding confidence to both retrieval models. In addition, we found that the effects of variable cloud opacity and temperature at altitudes less than 70 km were indistinguishable since the calculated extinction cross section of the assumed cloud particles has very little variation with wavelength across the 4 – 5 mm window. [14] To investigate whether the CO abundance in Venus’ mesosphere could indeed be retrieved from VIRTIS-M observations of Venus’ night side from 4 to 5 mm, a series of retrieval tests was first undertaken whereby a set of synthetic spectra were generated from a set of randomly varied atmospheres. Synthetic Gaussian noise was added to these spectra, whose amplitude was set by the expected NESR of VIRTIS for typical observations. These synthetic ‘measured’ spectra were then fed into NEMESIS to retrieve some or all of the quantities varied and the retrieved properties compared with the values originally used to
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Figure 2. Correlation of retrieved CO with synthetic (or ‘true’) CO for the tests where temperature and mean CO abundance were retrieved with NEMESIS from synthetic spectra generated by randomly varying the temperature, mean CO abundance, and cloud opacity and scale height about their a priori values. The agreement between retrieved and ‘true’ CO can be seen to be very good. Cases where the retrieved CO is significantly greater than the true CO are cases where the cloud scale height exceeded 7 km. generate the synthetic spectra. The error on the ‘measured’ spectra was taken to be the same NESR used to generate the random noise together with a degree of forward modeling error, which accounts for uncertainties in the line data and correlated-k approximation. [15] Analysis of the spectrum generated from our a priori atmosphere showed that we are sensitive to the abundance of CO just about the cloud tops, whose lines usually appear in absorption as temperature falls with height at this altitude for most regions of the planet. Unfortunately, we found that these observations could reveal very little about the vertical distribution of CO and were thus forced to assume a vertical profile and decided to assume that the mole fraction varies little with height in this region, retrieving a single scaling factor. [16] In the first test, the cloud profile was kept constant and a range of spectra generated by varying the temperature profile and CO mean mixing ratio. For the temperature profile, sine waves of temperature (varying with respect to height) were added to the a priori temperature profile with a random variation of phase, wavelength and amplitude. Similarly, the CO mean mixing ratio was varied randomly about the a priori value. NEMESIS then used the resulting synthetic spectra (to which Gaussian noise had been added) to retrieve the temperature profile and mean CO abundance, and the retrieved CO values were compared to the true values. In all cases the retrieved CO closely matched the ‘true’ CO values. However, we were concerned that variation in the cloud profile might be aliased as CO variations and so an additional set of test spectra was generated where
temperature, CO and the cloud profile (represented by two parameters: (1) the integrated visible optical depth from space to an altitude of 65 km and (2) a scale height above and below that altitude) were randomly varied. These test spectra were then used (1) to again retrieve the temperature profile and mean CO mole fraction and (2) to retrieve the temperature profile, mean CO mole fraction and the two cloud parameters. In both cases the correlation between the retrieved and true CO was found to be very good (Figure 2) and can be explained by the fact that variations in the cloud profile produced spectral effects that are indistinguishable from thermal variations and also that we are most sensitive to CO just above the clouds. The cases in Figure 2 where the fitted CO greatly exceeded the true value were where the cloud scale height exceeded 7 km, which is a value far greater than anticipated for Venus, especially near the poles, where values of less than 5 km are expected. Hence, when analyzing the real observations, discussed in the next section, the cloud profile was kept fixed at its a priori profile and only the vertical temperature profile and mean CO mole fraction retrieved.
4. Retrievals [17] Before analyzing the extended maps, the measurements in the VI0029 and VI0067 data sets, which had already been averaged into 20 20 pixel bins as described earlier, were further averaged by latitude into bins of width 10° and spaced every 5°, again to achieve Nyquist sampling. Since the Venus Express spacecraft was virtually over the south pole for both sets of observations, the emission
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Figure 3. Variation of the retrieved CO volume mixing ratio scaling factor (of an assumed profile with 100 ppm CO at all altitudes) with latitude for the VI0067 and VI0029 observations. The horizontal dashed line indicates the a priori scaling factor and the dash-dotted lines indicate the a priori error limits. The retrieved CO abundance is indicated by the solid line, with retrieved error limits indicated by the dotted lines. angle did not vary much with longitude and so the measured spectra could simply be averaged and a mean value of the emission angle assigned to each latitude bin. The a priori temperature profile was set as described earlier with an assumed temperature error of 10 K at all altitudes and a correlation length of 1.5 scale heights. For CO we initially assumed a scaling factor of 0.1 ± 0.2 of the a priori profile where the mole fraction was set to 100 ppm at all altitudes. To prevent the retrieved mole fraction ever going negative
NEMESIS was set to retrieve the logarithm of the CO scaling factor. [18] Figure 3 shows the variation of retrieved CO factor with latitude in the southern hemisphere for the VI0029 and VI0067 observations. It can be seen that in both cases at low latitudes the mean retrieved CO abundance is reasonably well constrained at a level of 40– 50 ppm with an error of about 15 ppm. However, poleward of 50°S, the retrieved abundance drops rapidly to around 10 ppm, although the
Figure 4. (left) Fitted temperature profile contour map, (middle) CO functional derivative (rate of change of mean radiance in the range 4 – 4.95 mm (mW cm2 sr1 mm1) with CO mole fraction) contour map, and (right) sections of the CO functional derivatives at 20°S (dotted line) and 70°S (solid line) for the combination of the VI0067 and VI0029 observations. 5 of 11
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Figure 5. Quality of fit to the VI0067 observations at 20°S and 70°S. Measured spectra are shown as points and the units of radiance are mW cm2 sr1 mm1. Fitted spectra are shown by the solid lines. The spectrum calculated for the same temperature profile, but with 0 ppm of CO is shown as the dotted line, while the spectrum calculated for the same temperature profile, but with 100 ppm of CO is shown as the dash-dotted line. The top row compares the spectra, while the bottom row shows the residual difference between the measured and fitted spectra. The small absorption spike at 4.69 mm, which remains even for the 0 ppm CO case, is due to a weak Q-branch absorption of CO2. retrieved error increases markedly, approaching its a priori value, indicating the solution is poorly constrained by the measurements. Intriguingly, while Figure 3 shows the low value of CO extending to the pole for the VI0067 observations, the retrieved abundance for the VI0029 observations seems to increase toward the pole. [19] To understand why the model has difficulty in retrieving a CO scaling factor in the polar collar around 70°S we need to look at the retrieved temperature profile, shown in Figure 4. From Figure 4 we can see that there is a marked temperature inversion in the 65– 70 km altitude region at these latitudes associated with Venus’ polar collar. Figure 4 also shows the CO functional derivative, that is to say the sensitivity of the calculated radiance (integrated over 4 – 4.95 mm) with respect to the CO abundance at all latitudes and altitudes, and sections are also shown at 20°S (typical midlatitude) and 70°S (in the middle of the polar collar). At 20°S we can see that the sensitivity to CO abundance is greatest at 68 km and that increasing the CO fraction at all altitudes reduces the radiance, indicating the
CO spectral lines to be in absorption. In contrast, at 70°S we can see that while the derivative of radiance with CO abundance is negative at some altitudes, it becomes positive in the temperature inversion region. Given that we can only retrieve a mean CO fraction from these measurements, having a situation where CO increases the mean radiance at some altitudes and reduces it at others leads to limited and ambiguous sensitivity to this species. As a test, we took the retrieved temperature profile and then recalculated the spectra assuming a CO abundance of (1) 0 ppm and (2) 100 ppm, shown in Figure 5, together with the spectrum calculated with the fitted CO value. In the region of the polar collar (70°S), very little difference between the spectra calculated with the retrieved CO abundance and the two trial abundances could be seen, whereas at midlatitudes (20°S) varying the mean CO abundance had a very clear effect on the spectra. Hence, while we can be reasonably confident of the retrieved CO abundance at low to mid latitudes, the abundance in the polar collar is effectively
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Figure 6. Measured radiance (in units of mW cm2 sr1 mm1) averaged over the range 4.8– 5.0 mm for the (left) VI0067 and (right) VI0029 observations (binned into 20 20 pixel bins and stepped every 10 pixels in both directions). unconstrained as is indicated by the retrieved errors differing little from the assumed a priori limits. [20] As a final test, we also attempted to fit the observations with a temperature profile alone, keeping the CO value fixed at 100 ppm. We found that allowing the CO abundance to vary led to significantly improved fits to the observations in regions where we are sensitive to CO, showing that CO indeed has a significant and measurable effect on the observed spectrum. [21] Having investigated the fitting of temperature and CO to the longitudinally averaged spectra, and determined at
which latitudes CO might be retrieved, we then took the mapped observations and attempted to retrieve maps of the mean CO abundance in the southern hemisphere. Figure 6 shows the measured radiance (integrated from 4.8 to 5.0 mm) for the two sets of observations, indicating the position of the southern polar vortex and Figure 7 shows the precision to which the spectra could be fitted with our model expressed n c2 1 X ððmi si Þ=ei Þ2 , where mi are the measured ¼ as n 1 n radiances of estimated error ei, si are the calculated synthetic radiances, and n is the number of observations.
Figure 7. Fitted c2/n for the (left) VI0067 and (right) VI0029 observations. 7 of 11
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Figure 8. Retrieved CO scaling factor for the (left) VI0067 and (right) VI0029 observations, where areas with c2/n > 3 have been masked (black areas). Note that a scaling factor of 1.0 corresponds to 100 ppm of CO. Areas where c2/n is less than 1 are regions where our model fits the observed spectra well and thus where we can have some confidence in our retrieved temperatures and CO abundances, if CO is detectable. Tests were conducted with a priori CO scaling factors of 1.0 ± 2.0, 0.1 ± 2.0 and 0.1 ± 0.2. In all cases very similar results were achieved since there is no trade-off to be made between vertical resolution and precision in these retrievals as the shape of the CO profile is fixed, resulting in a simple least squares solution for CO abundance that does
not depend on the initial guess. Here we will show the results for the case where the CO a priori scaling factor was set to 0.1 ± 0.2. Since we fit the log of the scaling factor, it is important to remember that the a priori error assumed by NEMESIS was actually the fractional error in the abundance, in this case 2.0. In regions where there is little CO information, the retrieved fractional error remains close to 2.0, while in regions where CO is better constrained the fractional error reduces to less than 0.5. Figure 8 shows the retrieved CO factor regardless of retrieved error, while
Figure 9. Fitted CO scaling factor for the (left) VI0067 and (right) VI0029 observations, where regions with a fractional error exceeding 0.4 and c2/n > 3 have been masked. 8 of 11
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Figure 10. Retrieved values of T(70 km) T(65 km) for the (left) VI0067 and (right) VI0029 observations, showing the inverted temperature profile at these altitudes over the polar collar. Figure 9 shows the retrieved CO factor for cases where the fractional error has been reduced to less than 0.4 and thus where we can be confident that the CO abundance really is being constrained by the observations. The large black areas in Figure 9, mostly coincide with regions of the polar collar temperature inversion around 70°S, as shown in Figure 10. [22] Comparing the retrieved CO maps of Figures 8 and 9 with the mean latitudinal variation for the two sets of observations shown earlier in Figure 3 we can see many
of the same features: roughly constant CO abundance at low to mid latitudes, decreasing in the polar collar (but with much less constraint) and then increasing again toward the pole in the case of the VI0029 observations. Both maps seem to indicate a higher abundance of CO near the evening terminator at 50°S – 70°S, although since the abundance of CO at similar latitudes on the morning terminator is not well constrained (as can be seen from Figure 9) owing to the vertical temperature structure there, it is not possible to determine if this is a real increase; it is possible that the CO
Figure 11. (left) Measured radiance (in units of mW cm2 sr1 mm1) averaged over the range 4.8 – 5.0 mm for the VI0038_00 observation, together with (right) the fitted CO abundance, again showing high CO over one of the two bright dipoles at 75°N, 0°E, indicative of rapid downwelling. 9 of 11
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abundance is higher at all latitudes between 50°S and 70°S, but only observable near the evening terminator. [23] The retrieved CO abundance near the pole for the VI0029 observation is particularly interesting as it seems to pick out high CO values over the polar dipole feature at 80°S, 45°E and 85°S, 45°W. This feature would appear to be absent in the VI0067 observation owing to the polar dipole being mostly on the day side in this case as indicated in Figure 6. To check if this was a repeatable feature, an additional observation, VI0038_00, which covers the south polar dipole, was analyzed in the same way to yield the fitted CO map shown in Figure 11. Again, we see high CO over one of the dipole hot spots at 75°N, 0°E, which would be consistent with this being a region of rapid downwelling dragging CO-rich air from higher altitudes to just above the cloud tops where it can be observed.
5. Discussion and Conclusions [24] Observations of the (1 – 0) absorption band of CO at 4.7 mm on Venus’ night side by the VIRTIS-M instrument of Venus Express have been used to investigate the spatial distribution of this gas above Venus’ cloud tops. Retrieval tests showed the effects of cloud top temperature and cloud opacity to be indistinguishable in the 4 – 5 mm range and that the retrieved CO abundance differed little if clouds were included or omitted. Hence, the cloud opacity was fixed and a vertical temperature profile fitted simultaneously with the CO abundance. [25] We find little spatial variation in the abundance of CO at midlatitudes, with a retrieved abundance of approximately 40 ± 10 ppm, a figure slightly higher than the levels observed by SPICAV/SOIR [Svedhem et al., 2007] of 10– 20 ppm at an altitude of 70 km. However, SPICAV/SOIR estimated the CO abundance at this altitude to decrease with height and as we are most sensitive to the abundance just above the cloud tops at 65 km, our results are broadly consistent. [26] Unfortunately, we find it very difficult to constrain the abundance of CO in the cold polar collar at 70°S as the retrieved temperature structure in the CO line-forming region masks the lines. However, the observations would suggest that CO increases toward the poles as we detect a significant signature of CO over Venus’ south pole, coinciding with the areas of the bright dipole feature, which would be consistent with rapid downwelling in these features bringing CO-rich air from high altitudes to just above the clouds where it can be detected. In addition, both observations suggest the possibility of higher abundances of CO between 70°S and 50°S near Venus’ evening terminator, although it is possible that this enhancement exists at all other longitudes, but cannot be detected owing to the CO lines being masked by the vertical temperature profile. [27] A caveat to the conclusion of high CO over the polar dipole is that the shape of the vertical profile of CO had to be assumed in these retrievals, owing to the observations having effectively no vertical resolution assumption in the line forming region (65 – 70 km). It was assumed that the abundance of CO did not vary with altitude, but if the CO abundance actually decreases rapidly with altitude and if cloud tops in the polar dipole are much deeper than elsewhere, as is thought to be the case, then we would see
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a longer path length of CO in this region due simply to CO’s vertical profile. However, since the polar dipole is thought to be a region of rapid downwelling and since we know the abundance of CO increases with height in the high atmosphere, where CO is formed by UV photolysis, it could be argued that this possibility is unlikely. To answer some of these questions and constrain the abundance of CO more closely will require more work and the analysis of many more observations than has been possible in this paper. One approach will be to analyze also the VIRTIS-H observations, which have much better spectral resolution, but poorer spatial sampling. This work is currently in progress, but the results are not yet mature enough to be presented here. In addition, limb observations would allow us to directly determine the vertical profile of CO above the cloud tops around the planet and help greatly to resolve any possible temperature/cloud ambiguities and we hope that such measurements can be conducted by VIRTIS during the remainder of its mission. [28] Acknowledgments. The Venus Express/VIRTIS team is acknowledged for their work. This research was made possible by funding from the United Kingdom Science and Technology Facilities Council.
References Bertaux, J.-L., et al. (2007), A warm layer in Venus’ cryosphere and high altitude measurements of HF, HCl, H2O and HDO, Nature, 450, 646 – 649, doi:10.1038/nature05974. Be´zard, B., C. de Bergh, D. Crisp, and J. Maillard (1990), The deep atmosphere of Venus revealed by high-resolution night side spectra, Nature, 345, 508 – 511, doi:10.1038/345508a0. Clancy, R. T., and D. O. Muhleman (1991), Long-term (1979 – 1990) changes in the thermal, dynamical and compositional structure of the Venus mesosphere as inferred from microwave spectral line observations of 12CO, 13CO, and C18O, Icarus, 89, 129 – 146, doi:10.1016/ 0019-1035(91)90093-9. Clancy, R. T., B. J. Sandor, and G. H. Moriarty-Schieven (2003), Observational definition of the Venus mesopause: Vertical structure, diurnal variation, and temporal instability, Icarus, 161, 1 – 16, doi:10.1016/ S0019-1035(02)00022-2. Collard, A., F. Taylor, S. Calcutt, R. Carlson, L. Kamp, and K. Baines (1993), Latitudinal distribution of carbon monoxide in the deep atmosphere of Venus, Planet. Space Sci., 41, 487 – 494, doi:10.1016/ 0032-0633(93)90031-V. Conrath, B. J., P. J. Gierasch, and E. A. Ustinov (1998), Thermal structure and para hydrogen fraction in the outer planets from Voyager IRIS measurements, Icarus, 135, 501 – 517, doi:10.1006/icar.1998.6000. de Bergh, C., V. I. Moroz, F. W. Taylor, D. Crisp, B. Bezard, and L. V. Zasova (2006), The composition of the atmosphere of Venus below 100 km altitude: An overview, Planet. Space Sci., 54, 1389 – 1397, doi:10.1016/j.pss.2006.04.020. Grassi, D., P. Drossart, G. Piccioni, N. I. Ignatiev, L. V. Zasova, A. Adriani, M. L. Moriconi, P. G. J. Irwin, A. Negrao, and A. Migliorini (2008), Retrieval of air temperature profiles in the Venusian Mesosphere from VIRTIS-M data: Description and validation of algorithms, J. Geophys. Res., doi:10.1029/2008JE003075, in press. Grinspoon, D. H., J. B. Pollack, B. R. Sitton, R. W. Carlson, L. W. Kamp, K. H. Baines, T. Encrenaz, and F. W. Taylor (1993), Probing Venus’s cloud structure with Galileo NIMS, Planet. Space Sci., 41, 515 – 542, doi:10.1016/0032-0633(93)90034-Y. Gurwell, M. A., D. O. Muhleman, K. P. Shah, G. L. Berge, D. J. Rudy, and A. W. Grossman (1995), Observations of the CO bulge on Venus and implications for mesospheric winds, Icarus, 115, 141 – 158, doi:10.1006/ icar.1995.1085. Hanel, R. A., B. J. Conrath, D. E. Jennings, and R. E. Samuelson (2003), Exploration of the Solar System by Infrared Remote Sensing, 2nd ed., Cambridge Univ. Press, Cambridge, U. K. Hong, Y., and B. Fegley (1997), Formation of carbonyl sulfide (OCS) from carbon monoxide and sulfur vapor and applications to Venus, Icarus, 130, 495 – 504, doi:10.1006/icar.1997.5824. Huebner, W. F., J. J. Keady, and S. P. Lyon (1992), Solar photo rates for planetary atmospheres and atmospheric pollutants, Astrophys. Space Sci., 195, 1 – 289, 291 – 294.
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Irwin, P. G. J., N. A. Teanby, R. de Kok, L. N. Fletcher, C. J. A. Howett, C. C. C. Tsang, C. F. Wilson, S. B. Calcutt, C. A. Nixon, and P. D. Parrish (2008), The NEMESIS planetary atmosphere radiative transfer and retrieval tool, J. Quant. Spectrosc. Radiat. Transf., 109, 1136 – 1150, doi:10.1016/j.jqsrt.2007.11.006. Krasnopolskii, V. A., and V. A. Parshev (1983), Photochemistry of the Venus atmosphere, in Venus, pp. 431 – 458, Univ. of Ariz. Press, Tucson. Lacis, A. A., and V. Oinas (1991), A description of the correlated-k distribution method for modeling nongray gaseous absorption, thermal emission, and multiple scattering in vertically inhomogeneous atmospheres, J. Geophys. Res., 96, 9027 – 9063, doi:10.1029/90JD01945. Mills, F. P., and M. Allen (2007), A review of selected issues concerning the chemistry in Venus’ middle atmosphere, Planet. Space Sci., 55, 1729 – 1740, doi:10.1016/j.pss.2007.01.012. Pollack, J., J. Dalton, D. Grinspoon, R. Wattson, R. Freedman, D. Crisp, D. Allen, B. Bezard, C. de Bergh, and L. Giver (1993), Near-infrared light from Venus’ night side: A spectroscopic analysis, Icarus, 103, 1 – 42, doi:10.1006/icar.1993.1055. Rodgers, C. D. (2000), Inverse Methods for Atmospheric Sounding: Theory and Practice, World Sci., Hackensack, N. J. Roos, M., P. Drossart, T. Encrenaz, E. Lellouch, B. Bezard, R. W. Carlson, K. H. Baines, L. W. Kamp, F. W. Taylor, and A. D. Collard (1993), The upper clouds of Venus: Determination of the scale height from NIMSGalileo infrared data, Planet. Space Sci., 41, 505 – 514, doi:10.1016/ 0032-0633(93)90033-X. Rothman, L. S., et al. (2005), The HITRAN 2004 molecular spectroscopic database, J. Quant. Spectrosc. Radiat. Transf., 96, 139 – 204, doi:10.1016/j.jqsrt.2004.10.008.
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Seiff, A. (1983), Thermal structure of the atmosphere of Venus, in Venus, pp. 215 – 279, Univ. of Ariz. Press, Tucson. Svedhem, H., D. V. Titov, F. W. Taylor, and O. Witasse (2007), Venus as a more Earth-like planet, Nature, 450, 629 – 632, doi:10.1038/nature06432. Tonkov, M., N. Filippov, V. Bertsev, J. Bouanich, V.-T. Nguyen, C. Brodbeck, J. Hartmann, C. Boulet, and F. Thibault (1996), Measurements and empirical modeling of pure CO2 absorption in the 2.3 mm region at room temperature: Far wings, allowed and collision-induced bands, Appl. Opt., 35, 4863 – 4870. von Zahn, U., S. Kumar, H. Niemann, and R. Prinn (1983), Composition of the Venus Atmosphere, in Venus, pp. 299 – 430, Univ. of Ariz. Press, Tucson.
R. de Kok, P. G. J. Irwin, F. W. Taylor, C. C. C. Tsang, and C. F. Wilson, Atmospheric, Oceanic, and Planetary Physics, Clarendon Laboratory, University of Oxford, Parks Road, Oxford OX1 3PU, UK. (
[email protected]) P. Drossart, LESIA, Observatoire de Paris, 5 place Jules Janssen, F-92195 Meudon, France. D. Grassi and A. Negra˜o, INAF-IFSI, Via del Fosso del Cavaliere, 100, I-00133 Rome, Italy. G. Piccioni, INAF-IASF, Via del Fosso del Cavaliere, 100, I-00133 Rome, Italy.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, E00B23, doi:10.1029/2008JE003140, 2008
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Composition of the Venus mesosphere measured by Solar Occultation at Infrared on board Venus Express A. C. Vandaele,1 M. De Mazie`re,1 R. Drummond,1 A. Mahieux,1 E. Neefs,1 V. Wilquet,1 O. Korablev,2 A. Fedorova,2 D. Belyaev,2 F. Montmessin,3,4 and J.-L. Bertaux3,4 Received 14 March 2008; revised 20 June 2008; accepted 4 September 2008; published 27 December 2008.
[1] Solar Occultation at Infrared (SOIR), which is a part of the Spectroscopy for
Investigation of Characteristics of the Atmosphere of Venus (SPICAV) instrument on board Venus Express, combines an echelle-grating spectrometer with an acoustooptical tunable filter. It performs solar occultation measurements in the IR region at a high spectral resolution better than all previously flown planetary spectrometers. The wavelength range probed allows for a detailed chemical inventory of the Venus atmosphere above the cloud layer, with an emphasis on the vertical distribution of the gases. A general description of the retrieval technique is given and is illustrated by some results obtained for CO2 and for a series of minor constituents, such as H2O, HDO, CO, HCl, and HF. Detection limits for previously undetected species will also be discussed. Citation: Vandaele, A. C., et al. (2008), Composition of the Venus mesosphere measured by Solar Occultation at Infrared on board Venus Express, J. Geophys. Res., 113, E00B23, doi:10.1029/2008JE003140.
1. Introduction [2] Venus is a very warm and dry planet with a dense atmosphere composed mainly of carbon dioxide (CO2, 96.5%) and Nitrogen (N2, 3.5%). Chemically active species, such as sulfuric bearing gases (OCS and SO2) and halides (HCl and HF) have already been reported (see de Bergh et al. [2006] for a general review on the composition of the atmosphere of Venus below 100 km altitude). Measurements have been performed essentially in the mesosphere below 100 km and below the clouds. Information about minor atmospheric constituents, their concentration, reactions, sources and sinks is incomplete, as for example only scarce measurements have been performed above 100 km altitude. In particular, photochemical models of the middle atmosphere would benefit from abundance measurements of Cl-bearing gases. [3] The Solar Occultation at Infrared (SOIR) spectrometer is an extension mounted on top of the Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV) instrument [Bertaux et al., 2007a]. SPICAV/ SOIR is one of the seven instruments on board Venus Express, a planetary mission of the European Space Agency (ESA) that was launched in November 2005 and inserted into orbit around Venus in April 2006 [Titov et al., 2006]. 1
Belgian Institute for Space Aeronomy, Brussels, Belgium. Space Research Institute, Moscow, Russia. 3 Service d’Ae´ronomie du CNRS, Verrie`res-le-Buisson, France. 4 Institut Pierre Simon Laplace, Universite´ de Versailles-Saint-Quentin, Saint Quentin en Yvelines, France. 2
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003140$09.00
[4] SOIR [Nevejans et al., 2006] is designed to measure at high resolution (0.15 cm1) the atmospheric transmission in the IR (2.2 – 4.3 mm) using solar occultations. This technique allows for the derivation of unique information about the vertical structure and composition of the Venus mesosphere. SOIR is the first high-resolution NIR spectrometer on board a spacecraft investigating the Venusian atmosphere and it enables a sensitive search for new minor species from the top of the clouds up to about 125 km of altitude.
2. Description of the Instrument [5] The instrument has already been extensively described elsewhere [Bertaux et al., 2007a; Mahieux et al., 2008; Nevejans et al., 2006] and will only be briefly described here. SOIR is an echelle-grating spectrometer operating in the IR, combined with an acoustooptic tunable filter (AOTF) for the selection of the diffraction-grating orders. The free spectral range (FSR) of the echelle spectrometer, i.e., the spectral interval in which there is no interference or superposition of light from adjacent orders equals 22.38 cm1, whereas the bandwidth of the AOTF was originally designed to be 20 cm1, as measured on ground before launch [Nevejans et al., 2006]. The real measured bandwidth of SOIR is 24 cm1 [Mahieux et al., 2008], creating some order leakage on the detector. The wave number domain that can be investigated by the SOIR instrument extends from 2256 to 4369 cm1, and is divided into 94 smaller ranges corresponding to the different orders (from 101 to 194). The detector width for orders 101 to 122 is smaller than the FSR of 22.38 cm1 and hence the detector will miss part of the spectrum. For orders 123 to
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Figure 1. Example of spectra obtained during one occultation (sunset 15 April 2007). Each transmittance is obtained by making the ratio of the solar spectrum seen through the Venus atmosphere to the unattenuated solar spectrum measured above the atmosphere. The selection of a spectral interval is achieved through the acoustooptic tunable (AOT) filter, tuned in this case to (a) 15809 kHz for diffraction order 121, (b) 19869 kHz for diffraction order 149, (c) 23031 kHz for diffraction order 171, and (d) 25742 kHz for diffraction order 190. In these particular ranges of wave number, the main absorption lines are from HDO (Figure 1a), CO2 (Figure 1b), H2O (Figure 1c), and CO (Figure 1d). 194 the inverse happens: the detector width is equal to or larger than the FSR and the detector will not be completely covered by the selected order. [6] The SOIR detector has 320 columns along the wave number axis and 256 rows along the spatial axis. The slit is projected on 32 fixed rows only. Since, owing to imposed telemetry limitations, only a data volume equivalent to 8 rows of 320 pixels can be retrieved per second, one is forced to bin the rows in eight groups of four rows, provided only one order (or AOTF frequency setting) is used during a given second. It is however possible to select up to four different orders (not necessarily sequential) per second, allowing us to gather a more versatile set of absorption lines. This reduces the maximum measurement time per order to 250 ms and implies that only 2 larger bins of 16 rows will be used if the complete slit height has to be covered. Later, the binning was changed to 2 bins of 12 rows because the outside rows of the illuminated part of the detector received a lower signal, because that part of the slit was too close to the edge of the Sun [Mahieux et al., 2008]. Background measurements are subtracted onboard from the measurements themselves. [7] Raw spectra, registered by SOIR and transmitted to Earth, need dedicated processing in order to upgrade them to a calibrated data set. This involves detector nonlinearity correction, spectral calibration and division by a reference solar spectrum. Ideally, the reference spectrum that is taken
outside the atmosphere would be measured with an identical relative slit position with respect to the solar disk. Attitude drift of the spacecraft, however, makes the slit float which results in a gradual linear change of the intensity. This effect is also corrected for Mahieux et al. [2008]. [8] A SOIR occultation observation can be taken either at sunset or sunrise. In the case of a sunset, the measurement cycle is started well before the instrument’s line of sight to the Sun intersects with the top layers of the atmosphere, and reference spectra are recorded (at a rate of 1 spectrum s – 1). Once the top of the atmosphere is reached, solar light is absorbed and the intensity of the recorded signal starts to decrease until the Sun gets so flattened by refraction that the spectrometer slit moves out of the refracted solar disk. One of the main advantages of solar occultations is that it is a self-calibrated technique in terms of transmission: dividing a spectrum obtained during the occultation by a reference solar spectrum recorded outside the atmosphere removes the solar signature and leaves a transmittance containing only information about the composition of the Venus atmosphere. The reference spectrum is in fact defined by selecting spectra recorded within the 40 s before the level at 220 km is reached. [9] Figure 1 gives an example of the evolution of the spectra through one occultation (sunset 15April 2007) in the orders 121, 149, 171, and 190 corresponding to the 2725– 2750, 3330 – 3357, 3820 – 3855, and 4245 – 4283 cm1
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Figure 2. Latitudinal distribution of measurements of (a) CO2, (b) CO, (c) HF, and (d) HCl. Dots represent measurements performed when Venus Express is separated from limb at 65 km tangent height by less than 5000 km and crosses when this distance is larger. In the latter case, the vertical resolution is coarser. Only measurements corresponding to the smallest distances have been considered in this study. Most of those correspond to north polar air masses. ranges, respectively. These transmittances show the characteristic behavior observed on all occultation series measured by SOIR. At the beginning of the series, the light path does not cross the atmosphere. No absorption signatures are present and transmittances are equal to unity. As the Sun
sets, the light path goes deeper and deeper into the atmosphere, and two absorption processes take place: the overall signal decreases owing to extinction by aerosols and absorption signatures appear. At the end of the observation, no light is captured anymore when the Sun disappears behind
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Figure 3. Typical examples of the vertical resolution for different types of orbits. The geometry of the orbit defines the field of view (instantaneous range of altitude encompassed by the instrument entrance slit taken at the beginning of the measurement, solid line) and the vertical resolution (dashed line). This corresponds to the portion of atmosphere sounded during one measurement, thus 250 ms. It highly depends on the distance from the planet and the velocity of Venus Express (VEX). Figures 3a and 3b show the vertical resolution and the field of view during orbits 223 (30 November 2006; distance to the limb at 65 km tangent height is 2000 km; 81°N) and 332 (19 March 2007; distance to the limb is 11198 km; 4°N) as a function of time. the cloud deck or moves out of SOIR’s field of view owing to diffraction. The structures seen in the spectra of Figure 1 are mainly attributed to HDO (Figure 1a), CO2 (Figure 1b), H2O (Figure 1c), and CO (Figure 1d). From Figure 1, it can also be seen that in general, the SOIR spectra contain information on the Venus atmosphere between 65 and 110 km for molecules such as HDO or HCl. For H2O and CO signatures are still observable up to 130 km altitude and CO2 features are seen up to 125 –130 km [Bertaux et al., 2007b; Wilquet et al., 2007]. [10] Most of the measurements of SOIR occur at high northern latitude because of the shape of the orbit with its pericenter located at about 250 km above the northern pole and its apocenter at about 65,000 km. When solar occultation occurs, a sunset or a sunrise can be observed. When the satellite is close to the planet the vertical resolution is less than 1 km. Measurements correspond to latitudes ranging from 60° to 90°N. When the satellite is located far from Venus, measurements have a poorer vertical resolution and occur at lower latitudes, typically from 70°S to 60°N. Figure 2 illustrates the latitudes and longitudes, corresponding to the tangent altitude of 65 km, of the measurements yielding information on CO2, CO, HCl, and HF. We have distinguished two types of geometries, when the distance between
Venus Express (VEX) to the limb (at 65 km tangent height) is less than 5000 km (dots), and when it is larger (crosses). Figure 2 will be discussed in more detail in section 4, where results for each species will be described individually. [11] We define the vertical resolution at the tangent point as the total altitude range scanned during one measurement. Because the slit is not always parallel to the limb, but rotating slightly, and because the measurement lasts for 250 ms, this vertical resolution may vary between 100 m and several kilometers in the worst cases. Another variable is the field of view, which is defined as the instantaneous height encompassed by the instrument slit at the beginning of the measurement. This parameter always gives a lower limit to the vertical resolution. The vertical resolution is mainly a function of the distance of the satellite to the planet. The distance of the spacecraft to the planet and the velocity of the spacecraft influence the vertical resolution in the following ways: (1) the further the spacecraft is from the limb, the larger the size of the instantaneous height measured in the atmosphere; (2) the velocity of the spacecraft projected at the limb on the atmospheric local vertical depends on the position of the spacecraft on its orbit around Venus at the moment of the occultation.
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Figure 4. Geometry of solar occultation measurements and definition of the onion- peeling method. [12] Figure 3 shows the variation of the field of view and vertical resolution as a function of time of orbit 223 (30 November 2006; distance to the limb at 65 km tangent height is 2000 km; 81°N) and 332 (19 March 2007; distance to the limb is 11,198 km; 4°N). During orbit 332, the spacecraft was far away from the planet which implies a large atmospheric height seen by the instrument slit. The field of view and vertical resolution are very similar because the Sun is rising almost vertically in the atmosphere and the velocity of the spacecraft is not too high (the spacecraft is almost at the equator). During orbit 223, spacecraft is close to the planet and the instantaneous height of the sounded atmosphere is much smaller. The velocity is higher, because the spacecraft is almost at pericenter (87°N). For the reasons explained just above, we will limit the retrieval to spectra obtained under favorable conditions, i.e., corresponding to low distance between VEX and the limb at 65 km tangent height (<5000 km) and with vertical resolution of less than 2 km. As can be inferred from Figure 2, this condition limits the selection to regions near the north pole, with latitudes above 60°N.
3. Description of the Retrieval Technique [13] The retrieval technique is based on the onion peeling method, illustrated in Figure 4. The atmosphere is treated as an onion-like composite of spherical layers, in which the temperature, pressure, and mixing ratios of the constituents are held constant.
model based on the fundamental knowledge of the optical properties of the atmosphere. In the forward model, SOIR spectra are simulated using a line-by-line (LBL) code developed initially for Earth exploration [Vandaele et al., 2006] and adapted for the conditions on Venus. The general equation describing radiative transfer through the atmosphere can be written as I ðn Þ ¼ I0 ðn Þetðn;0;sobs Þ þ
Z
Sobs
Bðn; T ðsÞÞaðn; sÞetðn;0;sÞ ds ð1Þ
0
where I0 represents the light intensity of the source (here the Sun) placed at the starting point of the raypath situated at the distance sobs from the observer, a(n, s) is the absorption coefficient, B(n, T) is the Planck function, and Z
s2
t ðn; s1 ; s2 Þ ¼
aðn; sÞds
ð2Þ
s1
is the optical depth along the path between the points s1 and s2. In solar occultation, the second term is often negligible with respect to the Sun intensity. The optical depths t of the absorbing constituents along the line of sight (see Figure 4) are calculated at high resolution in each layer considering the temperature and pressure determined by the ray-tracing procedure:
3.1. Forward Model [14] Quantitative analysis of the recorded spectra needs first the calculation of synthetic spectra through a forward 5 of 16
t ð8; n Þ ¼
Z að PðsÞ; T ðsÞ; n Þds s
ð3Þ
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where t is a function of wave number and depends on the tangent height. This, in turn, depends on the solar zenith angle 8 and on the atmospheric refraction characteristics, which, on Venus, are far from negligible. Starting from known temperature and pressure vertical profiles, the raytracing calculations are carried out on a finer grid (with a 200 m step) and the final results, i.e., the effective temperature and pressure in each layer as well as the effective densities, are obtained using the Curtis-Godson approximations [Goody and Yung, 1995]. In this work, we have used temperature and pressure vertical profiles from the VIRA model for altitudes up to 100 km [Seiff et al., 1985]. For higher altitudes (from 140 km and upward), data were taken from the model of Hedin et al. [1983] as suggested by Mueller-Wodarg and Tingle [2008]. The transition between the two data sets was performed by spline interpolating the temperature and reconstructing the pressure through the hydrostatic law. In the following we will refer to this composite model as the VENUSREF model. [15] The extinction coefficient a(P, T, n) is a function of the temperature and pressure prevailing at altitude z. It represents all absorption processes, including Rayleigh aR and aerosol aA extinction and absorption by molecular species: að Pð zÞ; T ð zÞ; n Þ ¼ aR ð z; n Þ þ aA ð z; n Þ þ
M X
si ð Pð zÞ; T ð zÞ; n ÞNi ð zÞ
ð4Þ
i¼1
where M is the number of absorbing species, si and Ni are the absorption cross section and number density of species i. The contribution of the aerosols will be further described and discussed in a later paper (V. Wilquet et al., Characterization of the upper Venusian haze from UV to mid-IR by SPICAV/SOIR on Venus Express, submitted to Journal of Geophysical Research, 2008). The contribution of each species is determined using a line-by-line model based on a line catalog specific to Venus, as explained later (see section 3.2). The absorption coefficient kij (cm2 molecule1) for a particular line j of species i is given as kij ðn; T ; P; pi Þ ¼ Sij ðT Þ f n; vij ; T ; P; pi
ð5Þ
where 8(. . .) is a normalized line shape. The intensity Sij(T) exhibits a temperature dependence which can be described as c2 n 0j T Qi ðT0 Þ c2 Ej T10 T1 1 e Sij ðT Þ ¼ Sij ðT0 Þ e c n 2 0j Qi ðT Þ 1e T0
ð6Þ
where c2 is the second Planck constant hc/kB (1.4387 cm K), with h the Planck constant, kB the Boltzmann constant, and c the speed of light, n 0,j is the central wave number of the jth transition, Ej is the energy of the lower state (cm1), and Sij(T0) is the intensity at the reference temperature T0 (cm1 molecule1 cm2). Q(T) and Q(T0) are the total partition functions under local thermodynamic equilibrium condi-
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tions, at temperature T and T0, respectively. These functions describe the temperature dependence of the line intensity of the transition. They are approximated by QðT Þ ¼ a0 þ a1 T þ a2 T 2 þ a3 T 3
ð7Þ
where a0, a1, a2, and a3 are tabulated coefficients [Gamache et al., 1990]. ASIMUT allows the user to select between different line profiles, the Voigt profile being the default. In the case of the H2O and CO2 molecules, sub- and superLorentzian line profiles have been observed [Clough et al., 1989; Pollack et al., 1993], which are characterized by a symmetric or asymmetric c function: fc ðn Þ ¼ cðn Þ fL ðn; P; T Þ
ð8Þ
This correction factor has been introduced to take into account the fact that far from the line center, the line displays marked deviations from the Lorentzian behavior 8L. Typically, CO2 displays a sub-Lorentzian behavior (the opacity far from the line center is less than that predicted by a Lorentzian profile) whereas H2O shows a superLorentzian behavior. The c function is usually defined on large spectral intervals. [16] The spectral grid, on which the profile must be determined, must be fine enough so that the narrowest line be adequately represented. In the upper atmosphere, the Doppler width is the limiting factor for selecting an adequate sampling value. As Doppler width depends on temperature, it varies with altitude; the line profile must therefore be sampled with different steps as the altitude varies. ASIMUT determines for each layer, of temperature T and pressure P, the optimized sampling step as DnðP; T Þ ¼
1 4
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi a2D ðT Þ þ a2L ð P; T Þ
ð9Þ
with aD the Doppler width obtained for a molecule of mass 20.0, and aL the Lorentzian width obtained for a molecule characterized by no self-broadening, a foreign broadening of 0.04 cm1 atm – 1, and a temperature coefficient of 0.5. However, for some particular temperature and pressure conditions, this sampling step might be too large compared to the desired final resolution. In that case, the sampling is set to the value of final resolution/15. LBL calculations often require the computation of a large number of line shapes over large to very large spectral intervals. The number of points might then become prohibitive. However, it can be observed that the line profile changes slower at a distance from the line center than it does near the center. The solution implemented in ASIMUT is the use of a nonuniform grid: near the center of the line, the optimized step derived with the help of equation (9) is used; the step is then progressively enlarged as one goes away from the center. Our algorithm is based on the study of Fomin [1995], which splits the spectral grid in a series of subintervals. Let us assume that the line shape must be calculated on the interval D, for which a value of 25 cm1 is considered sufficiently large for most of the simulations, except when applying a c factor. The cutoff value D is then determined by the interval on which this function is given.
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Figure 5. Description of the order overlap occurring in the instrument owing to the combined presence of the AOTF filter and the different diffraction orders of the echelle grating. In the top, a simulated spectrum is shown containing only CO2. It spans several diffraction orders whose limits are also indicated. This spectrum is first filtered by the AOTF, and only the portion under the filter bandwidth enters the spectrometer. Because this bandwidth is larger than the free spectral range (FSR) of the echelle spectrometer, more than one order of diffraction is transmitted. This is illustrated in the middle, where the contributions of the different orders are represented. Finally, in the bottom, the sum of all of these contributions is measured on the detector. The limits of the center zone are defined by [n - C, n + C] with C = 2/3 (aD + aL), aD and aL being the Doppler and Lorentzian widths of the line. The grid is divided into 2L portions located at unequal intervals:
n D; n 2L1 C ; . . . ; n 22 C; n 2C ; ½n 2C; n C ;
central zone; ½n þ C; n þ 2C ; n þ 2C; n þ 22 C ; . . . ;
ð10Þ n þ 2L1 C; n þ D
The number of intervals is related to D and C through the following expression: 2L C ¼ D
ð11Þ
In the central zone, the sampling is set to the optimized value determined with equation (9). Then in each subsequent interval, the sampling is doubled. This drastically reduces the number of points on which the line profile is calculated, without losing accuracy at the center of the line. At the end, the absorption line shape is interpolated to correspond to the sampling used for the determination of the optical depth (OD). The latter is chosen by considering the step asked for the radiance or transmission simulation and an oversampling factor either provided by the user or chosen such that the OD wave number step is a factor 10 lower than the final radiance step.
[17] Radiances are finally convolved by the instrumental function of the instrument, which is chosen to be a Gaussian, whose width varies between 0.13 and 0.25 cm1 depending on the spectral interval simulated [Mahieux et al., 2008]. [18] At this point, one must take into account the effect of the AOTF on the measured spectrum. As already noted, the filter has a bandwidth larger than the free spectral range of the echelle spectrometer, implying some order overlap on the detector. This is illustrated in Figure 5 where the contribution of four adjacent orders on top of the central one are considered. The incoming spectrum is represented in the Figure 5 (top), along with the position of the different orders of diffraction spanned by the AOTF filter bandwidth. The position of the maximum of the AOTF bandwidth is determined by the selected RF applied to the device. It is necessary to determine the best possible (AOTF frequency – wave number) calibration curve to increase the accuracy of the simulation. This is done through the careful analysis of specific measurements of some selected solar lines, whose positions are well known [Mahieux et al., 2008]. In the case illustrated here, the RF applied to the AOTF corresponds to the central order number 121. The entire signal received on the detector comes from the adjacent orders as no absorption is present in the incoming spectrum in the range of the central order of diffraction. The contribution coming from the noncentral orders is certainly not negligible and has to
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be taken into account when simulating SOIR spectra. In general, we consider a total of 7 contributing orders to determine the final transmittance. [19] The resulting transmittance is then interpolated to correspond to the wave number values of the observed spectrum. It is possible to fit a wave number shift between the observed and simulated transmittances to optimize the correspondence. 3.2. Spectroscopic Data Sets [20] Spectroscopic data, i.e., line parameters, have been taken from the latest version of the HITRAN database [Rothman et al., 2005]. However, broadening coefficients have been modified in order to take into account the presence of CO2 as main buffer gas, whereas data reported in HITRAN are given for Earth-like air conditions. [21] Sung and Varanasi [2005] reported CO2-broadened half widths and CO2-induced line shifts for the fundamental (1 – 0) of 12C16O at 201, 244, and 300 K, as well as for the first (2 – 0) overtone and the second (3 –0) overtone bands at 298 K. We have considered their values for the lines in common with the HITRAN database. Those which are not reported by Sung and Varanasi have been corrected by a conversion factor air width (shift) to CO2 width (shift) derived from the comparison of available common values. The temperature coefficient has been held constant for all lines (n = 0.73), as suggested by Sung and Varanasi. [22] Toth and Darnton [1974] have performed measurements of the HCl line widths by CO2 in the 1 – 0 and 2 – 0 bands of HCl. Their values have been taken for the lines in common with the HITRAN database. A correction similar to the one devised for CO was applied to the lines not reported by those authors. [23] Air-broadening measurements performed by the Brussels-Reims group [Fally et al., 2003; Jenouvrier et al., 2007; Me´rienne et al., 2003] up to 25,000 cm1 have revealed that there is a large vibrational dependence of the width for most of the H2O transitions. This is also the case under a CO2 rich atmosphere. Various attempts [Brown et al., 2007; Gamache et al., 1995] have tried to give simple relations between the quantum vibrational numbers and the width and shift of the lines. However, the spectrum of water is so complex that it is virtually impossible to obtain measured values for each band, reducing the validity of the proposed relations. A simplistic solution is to scale the broadening coefficients obtained with other perturbing gases, such as N2 or air, although it has been pointed out by Brown et al. [2007] that ‘‘simple scaling of existing values of air- or nitrogen-broadened parameters will not achieve sufficiently reliable CO2-broadened H2O coefficients’’, as these authors found considerable scatter of the ratio of CO2 to air broadened width (from 0.95 to 3.07) around the mean value of 1.67. However, no data, neither experimental nor calculated, are available in the spectral region investigated by SOIR and the constant factor of 1.67 was applied to correct the air-broadening coefficient given in HITRAN. This value has to be compared to the value of 1.3 usually used in Venus studies [Pollack et al., 1993]. 3.3. Onion-Peeling Method [24] The onion-peeling method was implemented to coherently treat a series of spectra recorded during one
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occultation. In this method, one starts the analysis in the uppermost layer, i.e., with the first spectrum containing absorption structures due to the constituents of the atmosphere, deriving concentrations in that layer, and progressively goes deeper into the atmosphere taking into account the results from the layers above. Vertical profiles of several key species of the Venusian atmosphere have been obtained by applying this technique, as will be demonstrated hereafter. [25] For the sake of clarity, we will consider in the following that the second right term in equation (1) is negligible. The observed transmittance Tr1 corresponding to ray 1 passing through the uppermost layer (layer 1 in Figure 4) is then given by
Tr1 ðn Þ ¼ exp a1 ðn ÞDs11
ð12Þ
where a1(n) stands for a(T1, P1, n) determined using equation (4) and Ds11 is the length of the raypath in layer 1 obtained by the ray-tracing procedure already explained. In this expression, the only unknowns are the concentration Ni of each species and the aerosol loading in layer 1, which are retrieved from the analysis of this first layer. Transmittance observed for ray 2 will result from the combination of the absorption of light in layer 1 (Ds21) and layer 2 (Ds22). If we moreover consider the atmosphere as spherical and homogeneous, we can further write
Tr2 ðn Þ ¼ exp a2 ðn ÞDs22 a1 ðn ÞDs21
ð13Þ
in which the only unknowns are the concentration Ni of each species and the aerosol loading in layer 2. By going down progressively, the vertical profiles of the interacting species can be derived.
4. Results [26] In this section, we will present some results of our analysis concerning the various species unambiguously detected with SOIR. CO2, being the most prominent absorber in the IR region, can be probed in different spectral regions. Combining different diffraction orders in one occultation moreover allows the possibility to use different bands presenting very distinctive absorption levels, therefore permitting the extraction of the CO2 vertical profile from the top of the clouds up to 170 km high. There is also the clear possibility to probe different absorption bands originating from the different isotopologues of the CO2. CO has been shown to act as a potential tracer of the dynamical processes occurring in the Venus atmosphere and its detection in the high-altitude range is of high interest. HCl and HF have also been detected in the SOIR spectra and vertical profiles are described. Some results will be briefly described concerning H2O and HDO but we refer to the paper of Fedorova et al. [2008] for a detailed discussion. Finally the detection limits of a series of species, which have not yet been unambiguously detected, are reported and discussed. 4.1. CO2 [27] CO2 is the main component of the atmosphere of Venus (96.5%). The first measurements of the atmospheric composition of Venus were made by Adams and Dunham [1932] using the 100 inch reflector at Mount Wilson. They
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Table 1. Possible Orders Where CO2 Can Be Detecteda Order
Isotopologue
Intensity
Temperature Dependence
101 102 103 104 105 106 107 108 109 111 112 115 116 117 118 140 141 142 143 147 148 149 150 151 154 155 156 157 158 159 160 161 162 163 164 165 166 167 168 169 170 171
628 628 626 (628) 626 (628) 626 (628) 626 (628) 626 626 626 628 628 628 628 628 628 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626 626
S S S S S S M W W W W W W W W W W W W W M M M W M M M M M M S S S M S S S S M W W W
M M S S S S W W W W W W W W W W W W W W W W W W W W W M M/W M M M M M M/W M/W M/W M M W W W
Category 1 1 1 1 1 1 2–3 3 3 3 3 3 3 3 3 3 3 3 3 3 2–3 2–3 2–3 3 2–3 2–3 2–3 1 2–3 1 1 1 1 1 2 2 2 1 1 3 3 3
a Indication on the altitude range probed and on the isotopologue measured. W, M, and S are for weak, medium and strong, respectively. Category is defined as the following: 1 means S/M intensity with S/M temperature dependence (cases for probing high altitudes CO2 and temperature); 2 means S/M intensity but W temperature dependence (cases for probing high altitudes CO2, less sensitive to temperature); 3 means M/W intensity with W temperature dependence (cases for probing lower altitudes CO2).
discovered three bands that they tentatively attributed to CO2. Since then, CO2 has been proven to be the main absorber in the infrared region. Its absorption bands are present throughout the spectral domain covered by SOIR, with intensities varying over a wide range of values. Combining different spectral intervals (orders of diffraction) in which the CO2 line strengths differ widely, the CO2 vertical profile can be obtained from lower altitudes around 65 km to higher altitudes of about 170 km. Indeed, the interval where the CO2 absorption is the largest will lead to information on the highest layers of the atmosphere, but will saturate for lower tangential height and, on the contrary, spectral intervals where the CO2 lines are weaker will provide information on the deepest layers. Moreover, the choice of the spectral interval has to be done with great care, because of the high temperature dependence of some
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absorption bands. This, in turn, could lead to the retrieval of temperature profiles if CO2 bands characterized by high temperature dependency are selected. In general, a mix of intervals is chosen such that one is optimal for the CO2 detection at high altitude, a second for CO2 detection at lower altitudes, and finally two intervals, with higher temperature dependency, are chosen to give potential information on the temperature as a function of altitude. A summary of the possible spectral intervals and their characteristics i.e., the altitudes sounded and their temperature dependence is given in Table 1. [28] The possibility to retrieve different isotopologues of CO2 has already been illustrated by the discovery of the 01111 – 00001 band of 12C16O18O in the SOIR spectra [Bertaux et al., 2007b; Wilquet et al., 2007]. Besides this new band, 12C16O18O abundances can be derived from well known features around 2500 cm1. A series of occultations dedicated to the measurement of the isotopologues which can be detected by our instrument were performed during MTP 21 (orbits 583 to 598) which occurred in January 2007. The spectral region in which SOIR is active contains a large number of lines owing to the three of the four main isotopologues of CO2. Spectral signatures of 12C16O16O, 12 16 18 C O O, and 12C16O17O have been clearly identified in the SOIR spectra, as shown in Figure 6. Simultaneous measurements of the different isotopologues will lead to the determination of the 17O/16O and 18O/16O isotopic ratios as a function of altitude. 4.2. HCL and HF [29] HCl and HF were observed for the first time in the Venus atmosphere by Connes et al. [1967], who estimated their mixing ratios at the cloud top as being 0.6 ppm (refined to 0.4 ± 0.07 ppm by Young [1972]) and 5 ppb, respectively. More recent nightside observations [Be´zard et al., 1990] provided measurements of HCl and HF below the clouds. These authors reported values of 0.5 ± 0.15 ppm and 5 ± 2 ppb, respectively, which are similar to the values found by Connes et al. [1967]. Preliminary measurements of Bjoraker et al. [1992] corresponding to altitudes above 72 km yielded a HF mixing ratio of 6.5 ± 0.3 ppb, in agreement with the values found previously. Recently, Iwagami et al. [2008] measured hemispheric distributions of HCl mixing ratio above and below the Venus cloud deck. These authors reported mean values of 0.76 ± 0.1 ppm at altitudes between 61 and 67 km and 0.4 ± 0.05 ppm at about 20 km. They argued that the larger HCl mixing ratio found above the clouds than that existing below the cloud requires the presence of a production process of HCl in the cloud region or above. [30] HCl abundances are retrieved from SOIR data using a series of lines belonging to the 1 – 0 band in the 2905– 2995 cm1 spectral range (orders 130 to 133), whereas HF determination is based on two lines of the 1 –0 transition (R1 and R2 at 4038.96 cm1 and 4075.29 cm1, respectively, in the orders 180 and 181). The measurement points reproduced in Figure 2 correspond to occultations where one of these orders was observed. Moreover, the P3 line at 3833.66 cm1 is also used to derive the HF abundance. This line, however, lies in a spectral region rich with H2O lines which render the determination of the HF abundance less accurate.
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Figure 6. The isotopologues 12C16O16O, 12C16O18O, and 12C16O17O have been observed during MTP 21 which was fully dedicated to the measurement of CO2 and the determination of the 17O/16O and 18 16 O/ O isotopic ratios as a function of altitude. Top shows observed spectra recorded by SOIR, and bottom presents the corresponding simulations. The asterisks correspond to 12C16O17O lines absorbing in the same region as 12C16O18O, and the circles correspond to 12C16O18O lines absorbing in the same region as 12C16O17O. [31] Figure 7 shows some typical vertical profiles found for HCl corresponding to the orbits 341 (28 March 2007; distance to the limb at 65 km tangent height is 3346 km; 82°N), 356 (12 April 2007; distance to the limb is 3298 km; 84°N), and 366 (22 April 2007; distance to the limb is 3898 km; 73°N). The actual parameters which are retrieved are the HCl densities in each of the sounded layers and volume mixing ratios are obtained by considering the air densities calculated from pressures and temperatures given by the VENUSREF model. The HCl profiles show very similar evolution with altitude. Almost systematically we observe a depletion feature around 90 km altitude. For comparison, a profile corresponding to a constant volume mixing ratio (vmr) (0.5 ppm) is also represented. The interpretation of the depletion feature observed around 90 km is not clear. An analysis with respect to temperature has been performed, which is illustrated in Figure 8, where we have considered four different vertical profiles for the temperature: the VENUSREF model, the VENUSREF model +20% and -20%, and a more realistic profile obtained by the Vera instrument [Pa¨tzold et al., 2007]. The four corresponding density profiles are compared in Figure 8: the absolute values change, because of the temperature changes, but the depletion feature does not disappear. This would indicate that the feature is not directly temperaturedependent. We have also investigated the sensitivity of the retrieval to inaccuracies in the determination of the position of the maximum of the band pass of the AOTF function. The results, also plotted in Figure 8, show that a displace-
ment of 1.0 cm1 of this maximum (the accuracy on this parameter has been estimated to be 0.83 cm1) [see Mahieux et al., 2008], implies a decrease of the HCl retrieved density by 10%, but the depletion feature does not disappear. [32] If we consider the evolution of the vmr with altitude of Figure 7, this corresponds to a more or less constant vmr comprised between 0.1 and 0.2 ppm, except around 90 km where it is slightly lower. The values at the lower boundary are somewhat lower than the values found in the literature, 0.4 ppm obtained by Connes et al. [1967] at an altitude of 64 km or 0.76 ppm at 61– 67 km from Iwagami et al. [2008]. Following the interpretation proposed by Iwagami et al. [2008] suggesting that there should be a chemical source of HCl inside or above the clouds to explain the gradient in HCl mixing ratio that they observed, our results seem to indicate that this potential source is restricted to the cloud region. [33] Profiles obtained in the case of HF are presented in Figure 9 for the occultations 357 (13 April 2007; distance to the limb is 3325 km; 83°N), 462 (27 July 2007; distance to the limb is 1864 km; 88°N), and 484 (18 August 2007; distance to the limb is 2670 km; 70°N). The HF abundance shows a more varying vertical distribution than HCl. Values found with SOIR seem to be in very good agreement with the values reported earlier, as for example, by Bjoraker et al. [1992], who measured HF abundances of 6.5 ± 0.3 ppb above 72 km. We note that both for HCl and HF, there is a
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Figure 7. Vertical profiles of the HCl density (left) and volume mixing ratio in ppm (right) obtained during the three occultations 341 (solid line, 28 March 2007; distance to the limb at 65 km tangent height is 3346 km; 82°N), 356 (dotted line, 12 April 2007; distance to the limb is 3298 km; 84°N), and 366 (dashed line, 22 April 2007; distance to the limb is 3898 km; 73°N). For comparison, the solid line with dots corresponds to a profile of constant volume mixing ratio (vmr) (0.5 ppm). The vmr in this plot correspond to the ratios of the HCl densities to the total air densities obtained from the temperatures and pressures of the VENUSREF model. Error bars correspond to the fitting error on the single-fitted parameter, i.e., the density in the corresponding layer, and does not take into account the errors on the densities in the upper layers. region of minimum vmr. However, their altitudes are different: 90 km for HCl, and 80 km for HF. 4.3. CO [34] The primary source of CO in the atmosphere of Venus is the photodissociation of CO2 by solar UV at altitudes higher than 120 km. The mixing ratio of CO was measured by the Pioneer Venus gas chromatograph [Oyama et al., 1980] at different altitudes in the lower atmosphere (between 22 and 62 ppm at 52 km, 30 ± 18 ppm at 42 km and 20 ± 3 ppm at 22 km). The gas chromatograph on Venera 12 [Gel’man et al., 1980] confirmed the low value found below 42 km (28 ± 7 ppm). Connes et al. [1968] reported a value of 45 ± 10 ppm at 64 km from Earth-based observations in the near-infrared. This value was corrected to 51 ± 1 ppm by Young [1972], who reanalyzed the spectra recorded by Connes et al. [1968]. These measurements seem to indicate the presence of a gradient in the mixing ratio of CO at least in the altitude range sounded. This was confirmed by observations of microwave lines of CO, which yielded CO mixing ratio for altitudes between 75 and 105 km. The CO mixing ratio increases from 55 ppmv at 75 km, to 130 ppmv at 85 km and 200 to 1000 ppmv at 105 km. Moreover it was shown that CO exhibits a significant diurnal variation but also strong year to year
variations [Clancy and Muhleman, 1991; Clancy et al., 2003; Gurwell et al., 1995]. Ground-based observations of the night side of Venus [Marcq et al., 2005; Marcq et al., 2006] have reported that the CO abundances in the lower atmosphere (below the clouds) showed a pronounced latitudinal enhancement of more than 10% when going toward the poles. Observations by the VIRTIS-M instrument on Venus Express [Irwin et al., 2008] have shown that there was little spatial distribution of CO just above the cloud (approximately 65– 70 km) at midlatitudes, with abundances of the order of 40 ± 10 ppm, with higher values at the poles, consistent with rapid downwelling bringing CO-rich air from higher altitudes. [35] Retrieval of the vertical CO profile from SOIR spectra relies on several absorption lines of the (2 – 0) CO band located between 4178 cm1 (order 187) and 4325 cm1 (order 192). All orbits during which this spectral region is investigated are shown in Figure 2b, distinguishing orbits corresponding to low and high (>5000 km) distance to the limb. Typically the vertical profiles are obtained from 70 to 125 km altitude, as can be seen from the examples reproduced in Figure 10. Those profiles correspond to the orbits 341 (28 March 2007; distance to the limb is3346 km; 82°N), 356 (12 April 2007; distance to the limb is 3298 km; 84°N), and 366 (22 April 2007; distance to the limb is 3898 km;
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Figure 8. Investigation of the temperature effect on the detection of HCl. In the left, the corresponding density profiles are shown. For comparison, the line with dots corresponds to a profile of constant vmr of 0.5 ppm. The feature around 85 km does not disappear. In the right, the relative difference in the retrieved HCl densities is shown for the different test cases. Four different temperature profiles have been considered: VENUSREF, VENUSREF with 20% and +20% excursions, and a vertical temperature profile obtained by the Vera instrument on board Venus Express [Pa¨tzold et al., 2007]. The effect of displacing the AOTF function maximum by 1.0 cm1 is also indicated. 73°N). CO vmr have been obtained from the retrieved CO densities using the air densities calculated from temperature and pressure from the VENUSREF model at the considered level. All three profiles show values for the CO vmr between 10 and 50 ppm below 90 km, and then increasing up to 4 – 8 104 ppm at 125 km, consistent with the existence of a source of CO, the photodissociation of CO2, at high altitudes. These profiles were obtained at the northern polar region and seem to be consistent with the observations performed by VIRTIS-M [Irwin et al., 2008] which indicate high values of the CO abundance at the south pole for altitudes between 65 and 70 km. They found values ranging from 100 to 200 ppm, these values being mostly influenced by the position relative to the brightest features of the polar vortex dipole. [36] Our value at 75 km (30 to 50 ppmv) is slightly lower than that derived from ground-based microwave (55 ppmv) by Clancy et al. [2003]. However, there is a major discrepancy above 75 km: our measurements show a decreasing vmr with increasing altitude, while microwave measurements indicate a steady increase. We find a minimum of CO (at 10 ppmv around 85 km). At higher altitudes, we find a steady increase of the vmr with altitude up to the limit of our measurements (4,000 – 8000 ppmv at 125 km). This steady exponential increase, when extrapolated upward to 140 km, would give a mixing ratio of 1, while in situ measurements with the Pioneer Venus Bus Mass Spectrometer yielded 0.4
[von Zahn et al., 1980]. The increase of vmr from 90 km upward is not due to diffusive separation of CO from CO2, because the homopause as determined by N2 measurements is at an altitude of 136 km. Rather, the CO gradient is dominated by the production rate of CO from CO2 which increases with altitude. More puzzling is the strong minimum region (80 – 90 km) found by SOIR, which had escaped detection up to now. It should be recognized that the technique of solar occultation provides an unprecedented vertical resolution. It could be of dynamic origin or it could be the result of a strong chemical sink. 4.4. H2O and HDO [37] Water is scarcely present today and it is not yet known if Venus was already dry at its formation or evolved slowly to its present state. In order to refine theories and models describing the formation and evolution of Venus and its atmosphere, to characterize the escape of D atoms from the upper atmosphere and to provide a robust explanation to the problem of the origin of water on Venus, high-resolution vertically resolved measurements of H2O and HDO and their temporal variations are needed. As illustrated in Figure 1, SOIR is able to measure both isotopologues of water simultaneously during the same occultation. Three absorption lines owing to HDO are clearly seen in the spectra recorded in the 2720 – 2725 cm1 region (order 121), whereas most of the features in the 3825 – 3855 cm1 region
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Figure 9. Vertical profiles of the HF density (left) and volume mixing ratio in ppb (right) obtained during the occultations 357 (solid line, 13 April 2007; distance to the limb at 65 km tangent height is 3325 km; 83°N), 462 (dotted line, 27 July 2007; distance to the limb is 1864 km; 88°N), and 484 (dashed line, 18 August 2007; distance to the limb is 2670 km; 70°N). For comparison, the solid line with dots corresponds to a profile of constant vmr (5 ppb). Error bars correspond to the fitting error on the single fitted parameter, i.e., the density (vmr) in the corresponding layer, and does not take into account the errors on the densities (vmr) in the upper layers. The following values from the literature are also indicated: Connes et al. [1967] (circles) and Bjoraker et al. [1992] (inverted triangles). (order 171) originate from the main isotopologue of water H2O and CO2. Such simultaneous measurements have been performed on a regular basis and are described in more detail by Fedorova et al. [2008]. 4.5. Other Minor Constituents [38] Spectra were recorded in different spectral intervals where some key constituents for the chemistry of the Venus atmosphere have prominent absorption features. However, some of those gases are only present as traces in the atmosphere and only upper limits of detection could be derived. As an example, OCS absorbs in the 2900–2950 cm1 range, corresponding to the order 130, but spectra recorded with SOIR contain no signature of this species, as illustrated in Figure 11. The upper limit of detection for OCS has been estimated to be 1.6 ± 2 ppb between 70 and 90 km and 0.02 ± 0.01 ppm above 90 km. OCS has up to now never been observed above the cloud deck. Measurements performed at lower altitudes [Marcq et al., 2005; Marcq et al., 2006] indicate that the OCS shows latitudinal variations with abundances between 5 and 20 ppm at 30 km decreasing with altitude. At 36 km, OCS amounts to 0.55 ± 0.15 ppm. The same spectral interval could also be used to derive the upper limit of detection of H2CO leading to the values of 3 ± 2 ppb below 90 km and 0.015 ± 0.01 ppm above 90 km.
[39] Similarly, it is also possible to determine the upper limit of other trace gases. In the case of SO2, the procedure is somewhat more complex since the SO2 signature is overlapped by a band of CO2, as explained and discussed by Belayev et al. [2008], who found a positive detection of 0.3 to 3 ppmv in some occultation profiles. 4.6. Error Analysis [40] The most critical parameters influencing the retrieval are related to the definition of the AOTF function: its formulation, the position of its maximum, and its extension in wave numbers (number of orders taken into account in the forward model). To test the sensitivity of our retrieval toward these parameters, several test cases have been investigated: (1) the function has been approximated by a sinc2 function or by a more complex function resulting of the sum of 7 sinc2, with different positions of their maxima and widths; (2) the number of adjacent orders has been varied between 2 and 4; (3) the position of the function maximum has been shifted by 1.0 cm1, a slightly higher value than the estimated accuracy on the determination of this position (0.83 cm1) [Mahieux et al., 2008]. The reference test case corresponds to the complex AOTF function using 3 adjacent orders (simulation on 7 orders in total), which corresponds to the normal settings used for retrieval. The results of this sensitivity study are the
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Figure 10. Vertical profiles of the CO density (left) and volume mixing ratio in ppm (right) obtained during the three occultations 341 (solid line, 28 March 2007; distance to the limb at 65 km tangent height is 3346 km; 82°N), 356 (dotted line, 12 April 2007; distance to the limb is 3298 km; 84°N), and 366 (dashed line 22 April 2007; distance to the limb is 3898 km; 73°N). For comparison, the solid line with dots corresponds to the CO in the VENUSREF model. Error bars correspond to the fitting error on the single fitted parameter, i.e., the density (vmr) in the corresponding layer, and does not take into account the errors on the densities (vmr) in the upper layers. following: (1) choosing a sinc2 function with 2 or 3 adjacent orders give rise to an underestimation of the CO2 density of 20 to 50% depending on the altitude, the maximum value
being obtained at around 90 km; (2) using the complex AOTF function but on only 2 adjacent orders implies an underestimation of 5 to 40%; (3) including more adjacent
Figure 11. SOIR spectra recorded around 2915 cm1 (order 130) where OCS and H2CO signatures should be observed. The main absorption structures observed in this region are due to HCl. Absorption features due to OCS and H2CO are also shown. From such spectra, only upper limits of detection can be derived. 14 of 16
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orders leads to higher densities by 1 – 2%, but lengthens the time required for the forward modeling; (4) changing the position of the AOTF maximum implies a difference of 5 to 20%. From this discussion, it is clear that the most critical factor in the simulation or retrieval of SOIR spectra is the accurate definition of the AOTF function, in particular its width and the positions and intensities of its sidelobes. In the near future, significant efforts will be devoted to a better yet determination of these parameters from specifically dedicated in-flight calibration measurements.
5. Conclusions [41] The SOIR spectrometer, which is part of the SPICAV/ SOIR instrument on board Venus Express has proven its high potential for the detection of minor key species for the understanding of the chemical and dynamical processes occurring in the Venus mesosphere. Detection of CO, HCl, HF, H2O, and HDO has been confirmed at altitudes ranging from 65 to 105 km, even 125 km depending on the species. Measurements have shown that the instrument was also sensitive to temperature through its observations of CO2 absorption lines, although this will require future development of the retrieval algorithm. SOIR is also able to differentiate between three isotopologues of CO2, namely 12 16 16 C O O, 12C16O18O, and 12C16O17O. This will in fine provide the vertical distribution of the 17O/16O and 18O/16O isotopic ratios. Vertical profiles of HCl, HF, and CO in the mesosphere above the northern pole show low variability in time. The CO vertical profiles measured by SOIR indicate values of CO abundances from 10 to 50 ppm at 90 km with a pronounced minimum of 10 ppm at altitudes of the order of 85 km, followed by an increase toward the higher altitudes. HCl volume mixing ratio is about 0.1 – 0.2 ppm and that of HF decreases from 1 to 7 ppb showing some variability from one orbit to the other. Because of its high sensitivity and its wide spectral coverage, SOIR is also a good tool to determine upper limit of detection for a series of trace gases: for example, OCS upper limit of detection has been estimated to be 1.6 ± 2 ppb below 90 km and 0.02 ± 0.01 ppm above 90 km. [42] Acknowledgments. The research program was supported by the Belgian Federal Science Policy Office and the European Space Agency (ESA, PRODEX program, contracts C 90268, 90113, and 17645). Procurement of AOTF was funded by CNES. Russian team also acknowledges RFBR grant 06-02-72563.
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from the surface to 100 kilometers altitude, Adv. Space Res., 5(11), 3 – 58, doi:10.1016/0273-1177(85)90197-8. Sung, K., and P. Varanasi (2005), CO2-broadened half-widths and CO2induced line shifts of 12C16O relevant to the atmospheric spectra of Venus and Mars, J. Quant. Spectrosc. Radiat. Transfer, 91, 319 – 322, doi:10.1016/j.jqsrt.2004.05.063. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017. Toth, R. A., and L. A. Darnton (1974), Linewidths of HCl broadened by CO2 and N2 and CO broadened by CO2, J. Mol. Spectrosc., 49, 100 – 105, doi:10.1016/0022-2852(74)90099-X. Vandaele, A. C., M. Kruglanski, and M. De Mazie`re (2006), Simulation and retrieval of atmospheric spectra using ASIMUT, paper presented at Atmospheric Science Conference, Eur. Space Agency, Frascati, Italy. von Zahn, U., K. H. Fricke, D. M. Hunten, D. Krankowsky, K. Mauersberger, and A. O. Nier (1980), The upper atmosphere of Venus during morning conditions, J. Geophys. Res., 85, 7829–7840, doi:10.1029/JA085iA13p07829. Wilquet, V., A. Mahieux, A. C. Vandaele, V. Perevalov, S. Tashkun, A. Fedorova, O. Korablev, F. Montmessin, R. Dahoo, and J.-L. Bertaux (2007), Line parameters for the 01111 – 00001 band of 12C16O18O from SOIR measurements of the Venus atmosphere, J. Quant. Spectrosc. Radiat. Transfer, 109, 895 – 905. Young, L. (1972), High resolution spectra of Venus: A review, Icarus, 17, 632 – 658, doi:10.1016/0019-1035(72)90029-2.
D. Belyaev, A. Fedorova, and O. Korablev, Space Research Institute, 84/32 Profsoyuznaya Street., 117997, Moscow, Russia. J.-L. Bertaux and F. Montmessin, Service d’Ae´ronomie du CNRS, BP3, 91371, Verrie`res-le-Buisson, France. R. Drummond, A. Mahieux, M. De Mazie`re, E. Neefs, A. C. Vandaele, and V. Wilquet, Belgian Institute for Space Aeronomy, 3 avenue Circulaire, B-1180 Brussels, Belgium. (
[email protected])
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HDO and H2O vertical distributions and isotopic ratio in the Venus mesosphere by Solar Occultation at Infrared spectrometer on board Venus Express A. Fedorova,1 O. Korablev,1 A.-C. Vandaele,2 J.-L. Bertaux,3,4 D. Belyaev,1 A. Mahieux,2 E. Neefs,2 W. V. Wilquet,2 R. Drummond,2 F. Montmessin,3,4 and E. Villard3,4 Received 18 March 2008; revised 13 June 2008; accepted 26 August 2008; published 25 December 2008.
[1] Vertical distributions of the molecular density and mixing ratios of H2O and HDO
in the Venus mesosphere have been obtained using Solar Occultation at Infrared (SOIR), a high-resolution (with l/dl 20,000) echelle spectrometer on Venus Express. The atmosphere is sounded in solar occultation in the range of altitudes from 65 to 130 km. Simultaneous measurements of water vapor lines in the spectral range around 2.61 mm (3830 cm1) at altitudes between 70 and 110 km and HDO lines around 3.58 mm (2715cm1) at altitudes 70–95 km have been performed. During 1 1/2 years, from April 2006 to August 2007, 54 such measurements have been carried out at different locations of Venus from the north pole to middle south latitudes. Most of the observations at morning and evening terminator correspond to high northern latitudes. We report values of mixing ratio and isotopic ratio obtained for 22 of those measurements occurring in the northern polar area. The average value of the volume mixing ratio of H2O is 1.16 ± 0.24 ppm and that of HDO is 0.086 ± 0.020 ppm. A depletion in the mixing ratio for both H2O and HDO is observed at 85 km, which can be related to a depletion of CO2 density above (95 km) and a possible temperature inversion at these altitudes. The vertical variation of HDO and H2O mixing ratio is within a factor of 2–3 for the analyzed set of observations. The temporal variations have been investigated, and no noticeable variability of H2O is reported at high northern altitudes. The average ratio of HDO/H2O obtained in this work, 240 ± 25 times the terrestrial ratio, is higher (1.5 times) than the value of 157 ± 30 times terrestrial reported for the lower atmosphere. This could be explained by a lower photodissociation of HDO and/or a lower escape rate of D atoms versus H atoms. Citation: Fedorova, A., et al. (2008), HDO and H2O vertical distributions and isotopic ratio in the Venus mesosphere by Solar Occultation at Infrared spectrometer on board Venus Express, J. Geophys. Res., 113, E00B22, doi:10.1029/2008JE003146.
1. Introduction [2] Venus’ atmosphere is significantly drier than the atmosphere of the Earth: it contains from 1 to 100 ppm of water depending on altitude and location. As reviewed by de Bergh et al. [2006], the water vapor content measured below the cloud level is within 20– 30 ppm, with much larger controversy at the cloud tops. [3] Most of the H2O mesospheric observations were obtained from ground-based millimeter heterodyne spectroscopy [Encrenaz et al., 1991, 1995; Sandor and Clancy, 2005]. Using the HDO line at 225.9 GHz and assuming a 1
Space Research Institute, Moscow, Russia. Belgian Institute for Space Aeronomy, Brussels, Belgium. 3 Service d’Ae´ronomie du CNRS, Verrie`res-le-Buisson, France. 4 Also at Institut Pierre Simon Laplace, Universite´ de Versailles Saint Quentin en Yvelines, Guyancourt, France. 2
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003146$09.00
D/H enrichment of 120 with respect to the terrestrial value, Encrenaz et al. [1991] obtained a mixing ratio of 3.5 ± 2.0 ppm at 60– 95 km. Later, observing the 187.31 GHz transition of H2O, Encrenaz et al. [1995] obtained 1 ppm above the clouds and 7 ppm (+5.0, 4.0) from the HDO line at 225.9 GHz, also assuming D/H = 120 terrestrial value. From a series of microwave observations, Sandor and Clancy [2005] reported a strong global variability of H2O on a 1 – 2 month time scale: at 65– 100 km the mixing ratio ranged from 0 ± 0.06 to 3.5 ± 0.3 ppm. Gurwell et al. [2007] also reported an extreme variability of the Venus mesosphere dramatically demonstrated by submillimeter wave astronomy satellite (SWAS) observations from December 2002. Over the course of 5 days, a deep ground-state water absorption feature consistent with a water abundance of 4.5 ± 1.5 ppm suddenly has transformed into a significantly shallower absorption, implying a decrease in the water abundance by a factor of 50 in less than 48 h.
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[4] The isotopic ratio of HDO to H2O in standard mean ocean water (SMOW) equals 3.1153 104 and corresponds to ([D]/[H])smow = 1.5576 104 [Hagemann et al., 1970]. On Venus, first measurements of the D/H ratio were performed by the Pioneer Venus mass spectrometer LNMS in the low atmosphere (100 (±12.5) times terrestrial, corrected in 1997 to 157 (±30) times terrestrial) [Donahue et al., 1982, 1997] and the ion mass spectrometer on the Pioneer Venus Orbiter at 155– 160 km ([D+]/[H+] = (1.7 ± 0.6) 102) or [D]/[H] = (2.2 ± 0.6) 102) at the turbopause level of about 132 km) [Hartle and Taylor, 1983]. These values have been confirmed later by ground-based observations in the 2.3 mm nightside windows [de Bergh C. et al., 1991]. Both in situ and remote-sensing determinations agree with one other and point to a strong deuterium enrichment in the Venus’ atmosphere. [5] In the near IR range above the clouds only one measurement of H2O and HDO has been reported [Bjoraker et al., 1992]. High-resolution spectra were obtained between 2.59 and 2.65 mm (3860 – 3770 cm1). A volume mixing ratio of H2O of 2.09 ± 0.15 ppm at 72 km was reported and a D/H ratio of 157 ± 15 times the value in the terrestrial ocean. Unfortunately, these results have never been published in any refereed literature. [6] The enrichment of 150 of deuterium to hydrogen supports the idea of dramatic escape of water from Venus. Two scenarios are possible: either Venus has had at least 0.3% of the terrestrial ocean and lost its water during a catastrophic process, or Venus was dry from the beginning, and the present isotopic ratio is explained by cometary impacts, degassing, and escape processes. The unknown vertical distribution of H2O in the middle and upper atmosphere precludes accurate theoretical descriptions of the escape of D and H atoms. New measurements of HDO and H2O are necessary to understand the evolution of the Venusian climate. [7] SOIR is a part of the SPICAV/SOIR experiment on the Venus Express spacecraft, operating on the orbit around Venus from April 2006 [Titov et al., 2006]. It is a high-resolution IR spectrometer working in the range of 2.2– 4.3 mm [Bertaux et al., 2007a; Nevejans et al., 2006]. The experiment is dedicated to vertical sounding of the Venus’ mesosphere in the range of 60– 120 km by means of solar occultation. One of the main scientific goals of SOIR is a simultaneous measurement of H2O and HDO vertical profiles to retrieve the isotopic ratio in the mesosphere. The first results on water vapor measurements with SOIR have been reported by Bertaux et al. [2007b]. In this paper we present a new analysis of the observed H2O content, employing a better instrument calibration and using longer series of observations.
2. Measurements 2.1. Instrument Description [8] Echelle spectrometer SOIR with an acoustooptic filtration of light is the first instrument with a spectral resolution above 20,000 orbiting another planet. It was first proposed for Venus Express mission by Korablev and Bertaux [2002], and built at the Belgian Institute for Space Aeronomy in collaboration with a Belgian industry in a very short time.
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[9] SOIR is designed to operate during solar occultations, when the instrument entrance optics is pointed toward the Sun as the latter goes down or up, allowing measurement of the atmospheric transmission of different layers of the atmosphere. The spectral range of SOIR from 2.32 to 4.25 mm (2353– 4310 cm1) allows measuring a number of atmospheric constituents showing up as absorbers in the transmission spectrum. High dispersion is provided by a 4 grooves mm1, arctan(2) incidence angle, echelle grating operating in diffraction orders from 101 (corresponding to the wavelengths of 4.4 mm) to 194 (2.3 mm). Each diffraction order covers a spectral interval from 20 to 40 cm1. The diffraction orders are separated by an acoustooptic tunable filter (AOTF). In the AOTF the radiation is filtered owing to volume diffraction on the acoustooptic wave excited within a birefringent crystal by a piezoelectric transducer at high frequency (RF = 14– 26MHz). The diffracted light is then analyzed by the echelle spectrometer. The central wavelength of the AOTF band-pass function is determined for the RF applied from the frequency-wavelength calibration as described by Mahieux et al. [2008]. The full width at half maximum FWHM of the AOTF band pass is around 24 cm1, and the profile of the AOTF band-pass function is close to a (sinx/x)2 function with a number of significant side lobes. As a result, spectral features leaking from several adjacent diffraction orders contribute to the spectrum observed on the detector with different weights that will be discussed in section 3. [10] The SOIR detector (from SOFRADIR) has 320 columns oriented along the spectral dispersion (wavelengths) and 256 rows along the spectrometer’s slit (spatial dimension). To avoid saturation, short integration times are used (20 to 30 ms), depending on the wavelength at which the measurement is taken. The background signal (Dark Current + Thermal emission of optics) is measured and subtracted onboard. In order to improve the signal-to-noise ratio (SNR), a number of measurements can be accumulated as long as the total measuring time remains below 250 ms. [11] The slit height is 30 arc min, and it is projected onto 32 rows of the detector. The slit width is equivalent to two detector pixels in the spectral direction. The spectral resolution of the spectrometer is very high and equals 0.13 cm1 at 2500 cm1 (order 111) and 0.27 cm1 at 4300 cm1 (order 192). The corresponding resolving power [l/Dl = n/Dn] is 20,000 [Mahieux et al., 2008]. [12] Owing to telemetry limitations, only eight spectra, each of 320 pixels long can be downloaded per second. During most of observations on Venus orbit, these eight spectra are taken in four different diffraction orders (different tunings of the AOTF), each corresponding to two large bins of 16 or 12 rows on the detector. [13] The detailed description of the instrument can be found by Nevejans et al. [2006] and Bertaux et al. [2007a]. Calibrations and in-flight performances of the instrument, including data handling, on board background subtraction, calibrations of the AOTF and the echelle spectrometer are described in detail by Mahieux et al. [2008]. 2.2. Observations [14] For accurate measurements of the isotopic ratio, simultaneous observations of H2O and HDO absorption lines are required. As discussed above, only four 20– 30 cm1
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Figure 1. Synthetic spectra of H2O (green), HDO (blue), and CO2 (red) at the target altitude of 85 km in solar occultation geometry. The top corresponds to the spectral interval 3760– 3940 cm1 (2.6 mm) dedicated to H2O measurements; the bottom shows the spectral interval 2660 –2840 cm1 (3.6 mm) dedicated to HDO measurements. Order boundaries are shown as solid (beginning) and dashed (end) black lines. The numbers above each plot show the order number and are located at the center of the order. Orders 171 and 121 are used mainly for retrieval of H2O and HDO abundance, respectively. portions of the spectrum within the spectral range of SOIR (four diffraction orders) can be acquired during a single occultation. We therefore carefully chose the spectral range for H2O and HDO detection attributing at least one order for HDO and another one for H2O. The priority was given to spectral ranges with strong lines of the gases of interest, and minimal contamination from other gases. There are several strong transitions of H2O within the spectral range of SOIR. The most preferable for the H2O retrieval is the strongest 3600 – 3900 cm1 band, which includes lines of n 1 and n 3 fundamental transitions. Below 3750 cm1 a strong CO2 2.7 mm band prevents the accurate determination of H2O, being completely saturated below 85 km of tangent altitude in solar occultation geometry. [15] For HDO we can consider the two strongest transitions within SOIR spectral range: the fundamental bands n 3 (001– 000) near 3707 cm1 (2.69 mm) and n 1 (100– 000) near 2723 cm1 (3.67 mm). The first one is blended by the 2.56 mm band of the main H2O isotope, and the HDO transitions are weak. The best candidate is the strong fundamental HDO band n 1 (100 – 000), completely isolated from H2O transitions. Moreover, this spectral range is free from detectable CO2 lines. Considering the other molecules, HCl lines are also located in the range of 3.67 mm. But these lines are isolated and do not contaminate HDO lines. [16] Synthetic spectra of Venus atmosphere limb transmission expected at the tangent altitude of 90 km, showing separately H2O and HDO lines, and all isotopologues of CO2 are presented in Figure 1 for the selected ranges
discussed above. The diffraction orders of SOIR are indicated. The orders numbers from 170 to 172 and 174 have been used for H2O measurements. Most of the observations so far have been performed in the order 171. The strongest lines of H2O in this range correspond to the rotational structures of (001) – (000) transition. For the retrieval of the HDO density, the orders 121 and 125 have been chosen, most of the observations having been performed in the order 121. 2.3. Observation Coverage [17] We have considered the operations from April 2006 to August 2007. During this period 54 orbits have been dedicated specifically to simultaneous measurements of H2O and HDO in the Venus’ mesosphere. Among them, on 47 orbits an additional spectral range dedicated to CO2 was scanned simultaneously in order to measure independently the atmospheric density profile and to reduce spacecraft pointing uncertainties. We address an interested reader to the paper by Vandaele et al. [2008, Figure 1] for the illustration of transmission evolution during an occultation in these diffraction orders. Vandaele et al.’s Figure 1 gives an example of the evolution of one occultation (sunset 15 April 2007) in the order 121, 171, 149, and 190. It is a typical observation dedicated to H2O, HDO, and CO measurements. Order 149 has been chosen for CO2 retrieval in the range of 3330 – 3357 cm1 where strong P and Q branches of 21102– 00001 transition allow the measurement of the atmospheric density at the altitudes of 70– 120 km.
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Figure 2. Spatial coverage for the simultaneous measurements of H2O and HDO. The maps of longitude versus latitude and local time versus latitude are presented. Most of the observations are located at high northern latitudes, corresponding to low distances from spacecraft to the limb. It reflects the peculiarity of the Venus Express polar orbit with the pericenter near the north pole. [18] The list of the main CO2 bands used for retrieval of atmospheric density is presented by Vandaele et al. [2008, Table 1]. In case of H2O and HDO, for the retrieval of the CO2 density several orders (depending on the orbit) have been chosen: 111, 112 (isotope O16C12O18, spectral range 2470 – 2520 cm1), 123 (spectral range 2748 – 2773 cm1) and 148, and 149 (a transition 21102 – 00001 at 3285 – 3360 cm1). [19] The geographical distribution of observations for which H2O and HDO were measured simultaneously, and the dependence on local time are presented in Figure 2. [20] Venus Express spacecraft is on highly elongated orbit with a pericenter located at about 250 km near the north pole, and an apocenter at 65000 km above the southern hemisphere of Venus [Titov et al., 2006]. From such an orbit, most of the solar occultations are observed at high northern latitudes with smaller distance to limb (near the pericenter), another part corresponding to relatively large distances to the limb and low latitudes. The vertical resolution of the occultation depends directly on the distance to the limb. The attitude of the spacecraft is maintained so that the slit of SOIR would be nearly parallel to the limb at 65 km. Near the pericenter the vertical resolution is generally better than 1.5 km. But because of the rotation of the slit with respect to the limb during one occultation and as the distance to limb is changing with the spacecraft motion, the vertical resolution during one observation usually varies from several hundred meters to several kilometers. This effect is described in detail in the paper of Vandaele et al. [2008, section 3]. In the present study we have considered only the measurements taken near the pericenter with small distances to the limb in order to minimize uncertainties. About 35 H2O measurements have been performed near the pericenter with the distance to the limb below 6000 km. All of them are located close to the north pole. We have further constrained the analyzed set to identical spectral ranges. We concentrated on orbits where the orders numbers 171 and 121 have been used for H2O and HDO respectively. In all, 22 orbits have been analyzed. The list of reported observations including universal time, latitude, longitude, local time, and distance to the limb at the altitude of 80 km is presented in Table 1. Within the considered set of observations which includes both sunrises
and sunsets, latitudes vary from 60° to 86° N, and the local time is 5 –6 h and 17– 19 h.
3. Retrieval Process [21] In solar occultation the transmission spectrum is directly obtained from relative measurements, and the photometric calibration of the instrument is generally not required. The processing of raw SOIR data, including nonlinearity and other corrections up to the construction of transmission spectra, and spectral calibration is described by Mahieux et al. [2008]. The reference spectrum is obtained as a weighted average of solar spectra recorded outside of the atmosphere (at the altitudes from 160 to 200 km) [Vandaele et al., 2008]. A forward model of gas absorption and the retrieval of vertical profiles of atmospheric constituents based on ‘‘onion-peeling’’ method are described in detail by Vandaele et al. [2008]. In the present paper we discuss only the details related specifically to H2O and HDO spectroscopy. 3.1. Spectroscopic Data [22] As a basis, the HITRAN 2004 database [Rothman et al., 2005] with the update of 2006 for H2O has been used. Nevertheless, no significant modification in the spectral range of interest was found in the 2006 update. Among the spectroscopic data like line strengths and positions, temperature coefficients and pressure shifts, HITRAN database contains line widths for air and self broadening. The widths of spectral lines broadened by CO2 which is the main component of the Venus’ atmosphere (96.3%) are not well known. The first measurements of CO2 broadening for H2O lines have been done by Howard et al. [1956a, 1956b, 1956c, 1956d]. This work recommended a constant factor of 1.3 to translate from the air to CO2 broadening. This factor has been then widely used in the 1990s for investigation of the Venus’ atmosphere [Pollack et al., 1993]. Later accurate laboratory measurements of CO2 broadening in the near-infrared range were made by Gamache et al. [1995]. The authors compared their laboratory measurements of CO2-broadened half widths with some recent experimental work and theoretical calcu-
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Table 1. Venus Express Orbits During Which Simultaneous Observations of H2O, HDO, and CO2 Have Been Performed With Low Distances to the Limb, Near the Pericenter of an Orbita Orbit
Date
Obs
Longitude
Latitude
LT
Ls
DTL
O1
O2
O3
O4
244 247 251 255 262 345 347 349 358 434 435 438 440 442 443 445 447 456 462 471 486 487
21 Dec 2006 24 Dec 2006 28 Dec 2006 1 Jan 2007 8 Jan 2007 1 Apr 2007 3 Apr 2007 5 Apr 2007 14 Apr 2007 29 Jun 2007 30 Jun 2007 3 Jul 2007 5 Jul 2007 7 Jul 2007 8 Jul 2007 10 Jul 2007 12 Jul 2007 21 Jul 2007 27 Jul 2007 5 Aug 2007 20 Aug 2007 21 Aug 2007
IN IN IN IN IN E E E E IN IN IN IN IN IN IN IN IN IN IN IN IN
248.91 256.82 261.13 268.06 282.80 13.69 36.86 81.05 194.10 264.07 266.74 274.52 279.49 284.24 286.51 290.81 294.66 298.02 254.94 232.25 251.33 253.47
84.80 84.39 82.09 79.19 70.76 85.93 87.34 88.13 82.30 70.83 72.54 76.11 77.84 79.30 79.97 81.18 82.31 86.52 87.68 79.49 66.41 63.01
19.96 19.88 19.34 18.99 18.54 7.56 8.69 11.23 16.90 5.88 5.85 5.76 5.68 5.59 5.53 5.41 5.26 3.64 23.55 18.98 18.42 18.36
63.08 67.82 74.14 80.46 91.52 224.28 227.52 230.76 245.36 7.99 9.59 14.36 17.54 20.72 22.31 25.49 28.67 42.93 52.39 76.02 90.16 91.78
2160.52 1558.68 1627.02 1729.67 2111.65 3315.49 3293.40 3281.31 3434.61 2060.91 1980.57 1848.39 1805.21 1777.70 1768.19 1758.40 1751.90 1780.27 1833.11 2088.11 2766.10 3001.76
121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121 121
171 171 171 171 171 171 171 171 171 171 171 171 171 171 171 171 171 171 171 172 171 171
111 149 112 148 111 149 149 149 149 111 112 149 118 149 111 149 111 112 149 123 149 149
112 192 181 191 133 133 133 190 133 125 190 132 151 184 188 130 136 145 180 180 190 190
H2O-1 (ppm) 1.08 ± 1.38 ± 1.00 ± 0.89 ± 1.02 ± 0.84 ± 0.84 ± 0.80 ± 0.75 ± 1.28 ± 1.14 ± 1.54 ± 0.96 ± 1.47 ± 1.22 ± 1.66 ± 1.22 ± 1.24 ± 1.60 ± 0.94 ± 1.54 ± 1.11 ±
0.10 0.13 0.10 0.05 0.06 0.06 0.04 0.05 0.05 0.11 0.03 0.10 0.09 0.05 0.07 0.10 0.05 0.04 0.09 0.12 0.10 0.05
H2O-2 (ppm) 1.24 1.36 1.28 1.03 1.08 0.68 0.74 0.63 0.64 1.15 1.21 1.50 1.25 1.35 1.20 1.53 1.22 1.26 1.59 1.36 1.31 0.99
± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±
0.13 0.17 0.05 0.05 0.07 0.06 0.04 0.04 0.05 0.12 0.03 0.15 0.05 0.07 0.07 0.14 0.04 0.02 0.13 0.04 0.18 0.05
HDOb (ppm) 0.097 ± 0.010 0.098 ± 0.012 0.092 ± 0.004 0.081 ±0.003 0.086 ± 0.004 0.056 ± 0.007 0.070 ± 0.014 0.054 ± 0.006 0.057 ± 0.011 0.075 ± 0.011 0.084 ± 0.002 0.106 ± 0.010 0.084 ± 0.003 0.091 ± 0.005 0.087 ± 0.004 0.106 ± 0.010 0.088 ± 0.006 0.086 ± 0.003 0.106 ± 0.011 0.118±0.004 0.081 ± 0.006 0.079 ± 0.005
a Latitude, longitude, and local time are given for the altitude of 80 km. Obs can be sunrise as E or sunset as IN; LT is the local time on Venus in Venusian hours; Ls is the solar longitude; DTL is the distance to the limb in km; O1, O2, O3, and O4 are the number of orders observed during the orbit; H2O-1 is the H2O mixing ratio in ppm averaged for altitudes between 75 and 112 km; H2O-2 is the H2O mixing ratio in ppm averaged for altitudes between 75 and 95 km; and HDO is the HDO mixing ratio in ppm averaged for altitudes between 75 and 95 km.
lations, and investigated the resulting temperature dependence of the half widths. The scaling factor from air broadening varies from 1.3 to 2.0, depending on a particular transition. Recently, Brown et al. [2007] published the CO2-broadened water parameters (half width, line shift, and temperature dependence of the widths) for the transitions between 200 and 900 cm1. They obtained ratios of CO2broadened to N2-broadened widths which varied widely from 0.95 to 3.07 with an average ratio of 1.67. To account for the CO2 broadening of lines in the Venusian atmosphere, we have multiplied the air-broadened half widths from the HITRAN 2004 database by 1.7 [Gamache et al., 1995; Brown et al., 2007]. The HDO line strengths in the HITRAN are scaled to the isotopic ratio of HDO to H2O in standard mean ocean water (SMOW) (3.1115 104). 3.2. Forward Modeling of H2O and HDO Transmittance Spectra [23] Using the forward model [Vandaele et al., 2008], the synthetic spectra of water vapor and HDO are calculated for the geometry of the solar occultation. Figures 3 and 4 illustrate the synthetic models for HDO in the order 121 (2703 – 2727 cm1) and for H2O in the order 171 (3823 – 3853 cm1) at a tangential altitude of 90 km. CO2 absorption lines have also been considered in those simulations. Atmospheric temperature and pressure profiles were taken from the Venus International Reference Atmosphere (VIRA) [Keating et al., 1985] for the dayside. H2O and HDO are assumed uniformly mixed with the volume mixing ratio of 1 ppm and 100 ppb, respectively. Monochromatic spectra have been converted to SOIR spectral resolution. Figures 3 and 4 (top) show modeling of gaseous absorption in ‘‘clear’’ case assuming there is no order mixing on detector. [24] The contribution of adjacent diffraction orders complicates the spectra measured by SOIR with respect to the
synthetic model shown in Figures 3 and 4 (top). The mixing of diffraction orders occurs owing to a wider than specified AOTF band-pass function (FWHM of 25 cm1), and the side lobes of this function (see discussion in the section 3.1 of Vandaele et al. [2008]). Figure 5 clearly demonstrates the contribution of different orders (actually ±3 orders) around the main order 121. Moreover, as described by Mahieux et al. [2008], the shape of the AOTF band-pass function is different for the beams coming from different parts of the spectrometer’s slit. As a result the contribution of adjacent orders is different for bins 1 and 2 recorded simultaneously during an observation (see section 2.1), where bin1 and bin 2 correspond to the top and bottom parts of the slit. [25] We reconstruct SOIR spectra in the 121 and 171 diffraction orders for bins 1 and 2 taking into account the overlapping of three orders from each side in Figures 3 and 4 (middle and bottom). In all, the seven diffraction orders contribute into the shown transmission spectra (from 3 to +3). [26] With the mixing of orders the 121 diffraction order, which is originally free from CO2 lines, becomes contaminated by CO2 lines coming from adjacent orders 122– 123 (compare Figure 1). This emphasizes the need of high accurate AOTF function taken into account more orders. [27] The mixing of diffraction orders becomes even more important in the order 171 (2.6 mm) located close to a strong 2.7 mm CO2 band. In Figure 5 we compare the synthetic model assuming the contribution of three diffraction orders from each side (Figure 5a) with a model assuming contribution of six orders (Figure 5b) with SOIR spectrum from the orbit 462 recorded at the altitude of 107 km. At this altitude the signal is not decreased by aerosol extinction, and the signal-to-noise of SOIR for order 171 between 100 and 290 pixels is better than 1000. It allows detecting very weak absorption features with a relative depth below 0.1–
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Figure 3. Synthetic models of HDO and CO2 in the order 121 (2703 – 2727 cm1). Frequency preset on the acoustooptic tunable filter (AOTF) is 15822.62 kHz. Solar occultation geometry is considered with the tangential altitude of 90 km. The VIRA atmospheric model was used. Volume mixing ration of HDO was assumed to be equal to 100 ppb. Monochromatic spectra have been converted to SOIR resolution. The thick solid line corresponds to HDO, and the thin solid line corresponds to CO2. H2O lines are also taken into account (dashed thin line) but are weak in this range. Top shows a ‘‘clear’’ spectrum of gaseous absorption in the range of 2703 – 2727 cm1 without any order mixing. Middle shows a synthetic spectrum taking into account the mixing of 7 orders (from 3 to +3) for bin 1 (spectra obtained in the bottom part of the slit). Bottom shows a synthetic spectrum taking into account the mixing of 7 orders (from 3 to +3) for bin 2 (spectra obtained in the top part of the slit). The different calibration of frequency-wavelength functions for different bins (or rows on the matrix) introduces differences in the spectra which correspond to different rows of the detector. 0.2%. Figure 5 demonstrates that the structures in order 171 are not a noise but CO2 lines coming from the orders 165 to 167. These orders correspond to the third side lobe of the AOTF function which has an amplitude of only 1% of the central lobe. However, the strong absorption in the CO2 2.7 mm band (up to 50%) results in a quasi-chaotic structure with relative depths of 0.1 – 0.3% inside the order 171, contaminating the measured H2O spectrum. [28] Although the AOTF band pass function is not well constrained for such distant side lobes, taking into account six adjacent orders allows to reduce the residual error. We therefore take into consideration the contribution from ±6 orders in all further modeling of H2O and HDO spectra in orders 121 and 171.
3.3. Fitting [29] The local densities of considered molecules (CO2, HDO, and H2O) are retrieved using the traditional onionpeeling method [e.g., Rodgers, 2000]. The main idea of this method in application to SOIR data is described by Vandaele et al. [2008] (see section 3.3 and Figure 4). The local densities of molecules are retrieved from the top of the atmosphere downward. The retrieved quantities are used for simulating the spectra corresponding to the lower layers. To calculate the absorption cross sections of the selected molecules, the atmospheric temperature-pressure profile was taken from VIRA [Keating et al., 1985]. [30] The measured spectra are fitted to the model using minimization of c2 and the simplex algorithm [Press et al., 1992] with two variables: local density of molecules (H2O,
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Figure 4. Synthetic models of H2O and CO2 in the order 171 (3823– 3853 cm1). The frequency preset on the AOTF is 22946.44 kHz. Solar occultation geometry is considered with the tangential altitude of 90 km. VIRA atmospheric model was used. Volume mixing ratios of H2O were assumed to be equal to 1 ppm. Monochromatic spectra have been converted to the SOIR resolution. The thick solid line corresponds to H2O, and the thin solid line corresponds to CO2. HDO lines are also taken into account (dashed thin line) but are weak in this range. The presentation of different top, middle, and bottom is the same as for Figure 3.
Figure 5. Contamination of adjacent orders in the main order 171 used for H2O retrieval. (a) ±3 orders around the main order have been taken into account. (b) ±6 orders around the main order have been taken into account. Those data were obtained during orbit 462 (27.07.2007) at the location latitude 87°N, longitude 245°E, 23.55 h Venusian local time (LT), distance to the limb 1830 km. 7 of 16
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Figure 6. Examples of data fitting for three orders 121 (HDO), 149 (CO2), and 171 (H2O) at the altitudes 86– 89 km. Those data were obtained during orbit 462 (27.07.2007) at the location latitude 87°N, longitude 245°E, 23.55 h LT, distance to the limb 1800 km. The light boxes relate to the spectral range chosen for HDO and H2O retrieval.
HDO, or CO2) and aerosol extinction that determines the continuum level of the transmittance spectrum. The best density value for every measurement is retrieved automatically, minimizing c2 statistically weighted according to the uncertainties of the measured quantity: c2 ¼
p320 X
ðTdata ðiÞ Tmodel ðiÞÞ2 =sðiÞ2
p1
where Tdata(i) and Tmodel(i) are the measured and modeled spectrum of transmittance for pixel i. s(i) is the statistical error determined from the signal-to-noise of the measured spectra which vary depending on diffraction order and detector pixel. Examples of the best fit spectra for H2O, HDO, and CO2 are presented in Figure 6. The light gray boxes in the top and bottom indicate the spectral range chosen for the retrieval of H2O and HDO density. The preference has been given to the lines with higher intensity, lower sensitivity to temperature, which are less contaminated with CO2 lines. Some lines appear broader in the model than in the observed spectrum that could be explained by the uncertainty of the assumed temperature profile (as partially described by Vandaele et al. [2008]), and possibly by sampling on the detector’s pixels. For the retrieval we used several lines that minimize the discrepancy in the line widths, as described in more detail in section 3.4. The curvature of measured spectrum in Figure 6 (top) is due to undercorrection of detector’s nonlinearity apparent at the edges of order’s spectral range. Figure 7 demonstrates a typical result of the retrieval of the gaseous densities for one
selected orbit, showing the simultaneous vertical distributions of the local densities for HDO, H2O, and CO2. 3.4. Uncertainty of Results [31] There are several sources of uncertainties in our results: spectroscopic data set (in particular the broadening factor for H2O in the CO2 atmosphere), atmospheric data set and assumed temperature-pressure profiles, and instrumental calibration uncertainties. The calibration uncertainties include AOTF frequency-wavelength calibration that determines the position of the maximum of the AOTF function versus the preset frequency for bin 1 and bin 2. This has been obtained with a precision of 0.83 cm1 [Mahieux et al., 2008]. Uncertainties on the shape of the AOTF function, including the contribution of the side lobes are still poorly estimated. Detailed sensitivity analysis on this parameter will be performed in the near future. [32] The uncertainties due to a shift of the maximum of the AOTF function of 1 cm1 give an error of 8% on the retrieved values at altitude from 70 to 95 km for HDO and 7% for H2O. The uncertainties due to the shape of the AOTF function (the cutoff of the far wing in ±3 orders during the retrieval that corresponds to removal of the contribution of the far side lobes) is the most important for H2O retrieval and put the error to 4% at the altitude from 70 to 100 km, which increases up to 40% at the altitude of 110 km. For HDO uncertainties from order mixing are below than 4%. [33] The uncertainties coming from the assumed spectral instrument function of the echelle spectrometer also give
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Figure 7. Example of retrieved densities of HDO, H2O, and CO2 in molecules cm3 for orbit 255 corresponding to the northern latitude 79.2°, longitude 261°E, and 19 h LT and distance to the limb 1730 km at altitude 80 km. rise to some systematic bias in the result. The sensitivity to variation of the instrument spectral function half width was investigated for each molecule. A variation of the half width by 20% results on average 8.5% of error in gaseous density, slightly varying with altitude and diffraction order. For different orders the widths of this function varies from 0.16 cm1 for the order 121 to 0.22 cm1 for the order 171 that may also bias the relation between different species detected in distant orders. [34] The comparison of results obtained for bin 1 and bin 2 also can provide a good estimation of uncertainties. For HDO (order 121) the total uncertainty on the average amounts to 8 – 14% and varies a little with altitude and from orbit to orbit that could be explained by different calibrations for these spectra and rotation of slit with respect to the limb. [35] Uncertainties on spacecraft pointing and in fact the unknown real pressure and temperature at the location of measurements add systematic biases to the retrieval. To account for this, we should use CO2 density retrieved from CO2 lines during the same observation as presented in Figure 7. The absorption cross sections of the molecules are sensitive to temperature and pressure. In particular, the line strengths depend on the temperature and the power of dependence is determined by the energy of transition (see equation 6 by Vandaele et al. [2008]). As already mentioned, in the case of H2O and HDO retrieval, CO2 density is retrieved from several different orders: 111, 112, 123, 148, and 149. The line strengths inside these ranges are not very sensitive to temperature variations. The intensity of the H2O and HDO lines are also not very sensitive to temperature, no more than 50% increase for a variation of temperature of 100 K. The uncertainties on temperature can be estimated by comparing the retrievals obtained with two different temperature profiles (Figure 8). We have chosen an extreme case that is assuming a deep inversion of temperature at the altitudes between 85 and 95 km. Retrievals of CO2 (order
149, top) and H2O (order 171, bottom) obtained for orbit 462 are presented in Figure 8. Left shows the temperature profiles used for the retrievals, middle shows the retrieved densities, and right shows the relative difference in % calculated as N1 ð zÞ N2 ð zÞ N1 ð zÞ
where N1(z) and N2(z) are the local densities for the first and second temperature profiles. For H2O, variations due to temperature are not very high and do not exceed 10%. For CO2, depending on the order, the error can reach 20% at the altitudes of the inversion (85 – 100 km). [36] The importance of the retrieved density profiles for the determination of the volume mixing ratios of the minor species is shown in Figure 9. Mixing ratios of H2O from orbit 442 obtained by division by the model density from VIRA or by the CO2 density retrieved in this work are compared.
4. Results [37] We report here results for 22 orbits obtained from January 2006 to August 2007. All the orbits have been performed at the latitudes of 63°– 88°N on morning and evening terminators with the distance to limb below 3500 km. [38] To detect variability of HDO and H2O abundance with time, several sets of observations have been done. In the present analysis the following sets are considered: 5 orbits from 244 to 262 at December 2006 to January 2007, 4 orbits from 345 to 358 at April 2007, and 15 orbits from 434 to 487 at June– August 2007. Figure 10 demonstrates the density retrieved for 5 consecutive orbits from 244 to 262 and 3 orbits from 442 to 445 obtained close to the north pole. No strong variations of the H2O and HDO
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Figure 8. Sensitivity of the retrieved parameters to the temperature profile. The results are presented for retrieval of CO2 (order 149, top) and H2O (order 171, bottom). They were obtained during orbit 462 (27 July 2007 at latitude 87.7°N, longitude 255°E, and 23.6 h LT, distance to the limb at altitude 80 km is around 1830 km). Left shows the temperature profiles used for the retrievals; middle shows the retrieved densities; right shows the residual error in % calculated as (N1(z)-N2(z))/N1(z), where N1(z) and N2(z) are the densities obtained for two different temperature profiles.
Figure 9. Comparison of the volume mixing ratios (vmr) of H2O using VIRA densities and CO2 densities derived from measurements recorded on the same orbit. The example shows the results obtained for orbit 442 (7 July 2007, at latitude 79°N, longitude 284°E, and 20.7 h LT, distance to the limb at altitude 80 km is around 1780 km). 10 of 16
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Figure 10. Vertical distributions of the H2O and HDO local density for eight orbits: five consecutive orbits from MTP09 (from 19 December 2006 to 10 January 2007) and three consecutive orbits from MTP016 (from 1 July 2007 to 27 July 2007), where MTP is the medium term plan. The conditions of observations are listed in Table 1. profiles have been detected for the altitude range of 75– 110 km. The latitude of observations varies for these orbits from 70° to 88°N. The mixing ratios of H2O and HDO for the same orbits are presented in Figure 11. [39] A marked depletion of H2O is observed in the range 80– 90 km, for which we have no explanation yet, other than noting that this altitude range coincides with the mesospheric minimum temperature and the top of the haze layer. The bump of H2O and HDO at 90– 100 km relates to depletion of CO2 density retrieved from CO2 order. Figure 9 demonstrates the difference between mixing ratios obtained assuming CO2 density from observations and from VIRA models. This depletion could be related to a strong temperature inversion with minimum at altitude 85– 90 km and we keep this study for a separate paper dedicated to the retrieval temperature and pressure profiles from CO2 lines. The curves of both isotopes show little variability. [40] The mean mixing ratio of H2O has been calculated for each orbit by averaging at altitudes from 75 to 112 km. The values are varying from 0.8 to 1.5 ppm. The temporal variation from orbit to orbit and the latitude variation of mixing ratios are presented in Figure 12. There is no systematic behavior in the distribution detected and the variations could reflect uncertainties of the SOIR observations. For HDO, mixing ratios were averaged from 75 to 95 km, the values varying from 50 to 130 ppb. Table 1 summarizes the averaged values of H2O and HDO volume mixing ratios (vmr). To make the comparison easier, the averaging of H2O for the same altitude range as HDO is
also presented in Table 1. The error bar corresponds to the error of the average in this case. [41] The vertical distribution of D/H ratio can be easily obtained from H2O and HDO density profiles. The complete set of results is presented in Figure 13. There are some vertical variations of the isotopic ratio that have been discussed by Bertaux et al. [2007b]. The average isotopic ratios scaled to the standard mean ocean water (SMOW) ratio of 3.10693 104 and averaged at the range of altitudes from 75 to 90 km are shown in Figure 14 as a function of orbit number and latitude. The obtained values range between 200 and 300 times terrestrial with the average of 240 ± 25.
5. Discussion [42] There were no continuous sets of H2O observations in the Venusian mesosphere before Venus Express. Sandor and Clancy [2005] and Gurwell et al. [2007] have made systematic observations of the mesosphere from groundbased and orbital telescopes. The summary of previous measurements is presented in Table 2. Most of the groundbased observations correspond to measurements of the Venusian disc and the direct comparison with the local observations of SOIR is not possible. Ground-based observations cannot directly resolve the vertical distribution of the species and inversion techniques are required to derive the profile from a single observation. The advantage of SOIR is that direct measurements of vertical profiles from the orbit are made possible. Moreover, simultaneous obser-
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Figure 11. Vertical distributions of the H2O and HDO volume mixing ratios for eight orbits: four consecutive orbits from MTP09 and four consecutive orbits from MTP016. The conditions of observations are listed in Table 1. vations in spectral ranges corresponding to both H2O and HDO are sometimes difficult to realize from the Earth. Several observations of H2O [Encrenaz et al., 1991; Sandor and Clancy, 2005] are, in fact, measurements of HDO lines and rely on the HDO/H2O ratio obtained from other measurements. The volume mixing ratio of H2O (1 ppm)
measured by SOIR is well consistent or lower than obtained from the Earth. The HDO mixing ratio of 0.08 ppm from SOIR is also lower than values obtained by Encrenaz et al. [1991, 1995] and by Sandor and Clancy [2005], but all values are in agreement within the error bars.
Figure 12. Temporal and latitudinal evolution of the H2O volume mixing ratios averaged as described in the text for orbits listed in Table 1. The H2O mixing ratio has been obtained by division by CO2 densities from the same observations. 12 of 16
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Figure 13. Vertical distributions of the HDO/H2O ratio for several orbits. [43] High variability of H2O in the Venus mesosphere was recently detected by Sandor and Clancy [2005] and Gurwell et al. [2007]. This variability is not supported by SOIR observations at least near the north pole region. The
SOIR values show variations, but no more than a factor of 2 around the mean value (1 ppm), staying inside the SOIR uncertainties (except perhaps in the lowermost range 70– 75 km, where errors are greater). But the direct compar-
Figure 14. Temporal and latitudinal evolution of the HDO/H2O ratio scaled to Earth’s value. 13 of 16
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Table 2. Comparison of Averaged Values of the H2O and HDO Mixing Ratios With Previous Observations Altitude (km)
Spectral Range
H2O (ppm)
HDO (ppm)
Reference
1990
60 – 95
225.9GHz (HDO)
0.13
Encrenaz et al. [1991]
1990 – 1993
60 – 95
187.31GHz (H2O) 225.9GHz (HDO) 2.59 – 2.65 mm (HDO and H2O) 226 GHz (HDO) and 335 GHz 547.676GHz (H2O) and 556.936GHz 2.61 mm (H2O) 3.58 mm (HDO)
3.5 ± 2 (D/H = 120 terrestrial value) 1 ppm
0.26
Encrenaz et al. [1995]
2.09 ± 0.15
0.102
Bjoraker et al. [1992]
0 ± 0.06 3.5 ± 0.3
0 0.17
Sandor and Clancy [2005]
4.5 ± 1.5
-
Gurwell et al. [2007]
0.8 – 1.5
0.05 – 0.13
This research
Years
1992?
72
1998 – 2004
65 – 100
2002 – 2004
65 – 100
2006 – 2007
75 – 95
ison with the ground-based observations is not possible. SOIR measures only morning or evening profiles at terminator, whereas values of H2O from millimeter observation may reflect disc averaged variability. It is equally difficult to compare our results to the H2O afternoon bulge observed by Pioneer Venus orbiter at the altitudes of 62– 70 km in low and middle latitudes from 10°S to 60°N [Koukouli et al., 2005] Therefore, more SOIR occultations in wider latitude range are required to understand spatial variations of water vapor in the mesosphere. [44] Several measurements of the D/H ratio in the Venus’ mesosphere have already been performed before. The 2.3 mm window has been used to observe lines from both H2O and the deuterated water HDO and to measure the D/H ratio below the clouds from ground-based observations. de Bergh et al. [1991] derived a value of 120 ± 40 times the terrestrial SMOW value from high-resolution CFHT spectra on the nightside. Donahue et al. [1982, 1997] has reported the first determinations of the D/H ratio performed in situ by the neutral mass spectrometer of Pioneer Venus, the analysis of their data yields a value of D/H = 100 ±12.5 corrected later to 157 ± 30 times the ratio in Earth’s oceans. Bjoraker et al. [1992] have obtained the value of 157 ± 15 times terrestrial using near-IR dayside ground-based observations. The isotopic ratio retrieved from SOIR measurements (240 ± 25) times terrestrial is higher (factor 1.5) than the results of all previous measurements in the lower atmosphere 157 ± 30 [Donahue et al., 1997] and the measurements by Bjoraker et al. [1992] at the effective altitude of 72 km but they support the higher abundance of HDO to H2O compared to Earth. [45] One motivation to HDO and H2O measurements in the upper atmosphere of Venus was to determine if HDO was present up to the altitude of photodissociation, providing D atoms for further escape. Clearly this D escape needs to be quantified, if we want to extrapolate back in time what was the original content of water on Venus. In the Earth’s stratosphere, there is a cold trap at the tropopause, and preferential condensation of HDO versus H2O is an important factor of fractionation, which is circumvented by deuterated methane (CH3D) passing through the tropopause without condensing [Moyer et al., 1996]. On Mars, the observed depletion of atomic D in the upper atmosphere [Krasnopolsky et al., 1998] was explained by the lower HDO photolysis rate [Cheng et al., 1999],
preferential condensation of HDO [Fouchet and Lellouch, 2000; Bertaux and Montmessin, 2001], and the smaller observed abundance of H2 [Krasnopolsky and Feldman, 2001] than expected from models. Thermal and nonthermal escapes of D are weaker than those of H [Krasnopolsky, 2002] and tend to increase D/H. [46] Our present results show that on Venus, there is no condensation of H2O nor HDO, and no cold trap preventing HDO to be photodissociated in the region above 80 km. On the contrary, one has to explain why the ratio HDO/H2O is found 1.5 times higher than lower in the atmosphere, with a possible trend of increasing with altitude (Figure 13). Bertaux et al. [2007b] proposed two explanations: the higher photodissociation rate of H2O [Cheng et al., 1999] will preserve more HDO; also, if D atoms are not at all escaping at the top of the atmosphere, they may eventually recombine with OH radicals, generating a downward HDO flow, and further decrease the importance of HDO photodissociation. In any case, detailed modeling of these mechanisms is necessary, but it is beyond the scope of the present paper.
6. Conclusions [47] We report vertical distributions of the molecular density and mixing ratios of H2O and HDO in the Venus mesosphere by SOIR, a high-resolution (with R20,000) echelle spectrometer on board Venus Express. The spectrometer operates in solar occultation mode and sounds the atmosphere at the range of altitudes from 65 to 130 km. Simultaneous measurements of water vapor lines in the 2.61 mm range (3830 cm1) at altitudes of 70– 110 km and HDO lines in the 3.58 mm range (2715cm1) at altitudes of 70– 95 km have been performed. For a year and a half, from April 2006 to August 2007, 54 such measurements have been carried out at different locations on Venus from the north pole to middle south latitudes mainly in the high northern latitudes at the morning and evening terminator. After in-flight recalibration of the spectrometer during summer 2007, a new analysis of the observed H2O has been undertaken. We analyzed the mixing ratio and isotopic values obtained from 22 orbits corresponding to high northern latitudes. The temporal variations have been investigated. The averaged volume mixing ratios of H2O = 1.16 ± 0.24 ppm and HDO = 0.086 ± 0.020 ppm have been obtained for this set of orbits. The depletion in the mixing ratios of H2O and HDO near 85 km is related to the
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depletion in the CO2 density near 95 km and a possible temperature inversion at these altitudes. No large variability of H2O has been detected for the high northern latitudes observations. The time variation of the mixing ratios of HDO and H2O do not exceed 2 –3 times the mean value. The obtained HDO/H2O ratio equals 240 ± 25 times the ratio in the Earth’ ocean on the average, which is higher than previously reported values by a factor 1.5. [48] Future analysis of SOIR data corresponding to middle and lower latitudes of Venus and to a wider set of observations will allow understanding the possible long-term variations. We also count on more accurate instrumental calibration and improved simultaneously retrieval of temperature from SOIR data. [49] The resulting vertical profiles of water vapor and its isotopologue could give an exciting challenge to the development of new photochemical models of the upper atmosphere of Venus. Such investigations have already started with models such the one developed by M.-C. Liang and Y. L. Yung (Modeling the distribution of H2O and HDO in the upper atmosphere of Venus, submitted to Journal of Geophysical Research, 2008). Moreover, combination of VIRTIS and SPICAV nadir measurements of H2O below and above the clouds could reconstruct a global distribution of water vapor in the Venus atmosphere from the lower atmosphere to the upper mesosphere [Drossart et al., 2007; Bertaux et al., 2007a; Marcq et al., 2008]. [50] These profiles give also hints to estimation of past and present escape from the planet. If there were no escape of D atoms, now and in the past, the present 1 cm (equivalent liquid) and D/H ratio 0.025 (enrichment 150) would imply in the past only 1.5 m, compared to 2.8 km on Earth. With the present observation of plenty of HDO in the photodissociation region, D atoms are certainly present in the thermosphere. It would be important to quantify the escape of D atoms, possibly by measuring mass 2 ions escape as could do ASPERA on board Venus Express [Barabash et al., 2007]. In spite of the possible confusion with H+2 ions, it would provide a useful upper limit to D+ ion escape [McElroy et al., 1982; Hartle and Taylor, 1983]. [51] Acknowledgments. We would like to thank our reviewers for helpful comments that improved the manuscript. We thank our collaborators at the three institutes for the design and fabrication of the SOIR instrument, which was mainly built in Belgium by OIP company, under the direction of IASB-BIRA. Russian team acknowledges RFBR grant 06-0272563. Belgium team was supported by the Belgian Federal Science Policy Office and the European Space Agency (ESA, PRODEX program, contracts C 90268, 90113, and 17645). Procurement of AOTF was funded by CNES and the French authors are sponsored by CNRS and CNES. We thank A. Bensoussan (COB) for procuring in due time the SOFRADIR detector.
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D. Belyaev, A. Fedorova, and O. Korablev, Space Research Institute, 117997, 84/32 Profsoyuznaya Street, Moscow, Russia. (
[email protected]. rssi.ru) J.-L. Bertaux, F. Montmessin, and E. Villard, Service d’Ae´ronomie du CNRS, BP 3, F-91371, Verrie`res-le-Buisson, France. R. Drummond, A. Mahieux, E. Neefs, A.-C. Vandaele, and W. V. Wilquet, Belgian Institute for Space Aeronomy, 3 avenue Circulaire, B-1180 Brussels, Belgium.
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First observations of SO2 above Venus’ clouds by means of Solar Occultation in the Infrared D. Belyaev,1 O. Korablev,1 A. Fedorova,1 J.-L. Bertaux,2,3 A.-C. Vandaele,4 F. Montmessin,2,3 A. Mahieux,4 V. Wilquet,4 and R. Drummond4 Received 18 March 2008; revised 29 July 2008; accepted 2 September 2008; published 31 December 2008.
[1] Solar Occultation in the Infrared (SOIR) is a part of the Spectroscopy for Investigation
of Characteristics of the Atmosphere of Venus (SPICAV)/SOIR occultation experiment on board Venus Express dedicated to the study of gaseous and aerosol vertical structure of Venus’ mesosphere. SOIR is an echelle spectrometer with acoustooptic selection of diffraction orders operating in the wavelengths range of 2.2–4.3 mm at high spectral resolution (l/Dl 20,000). Detection of minor constituents such as CO, H2O, HDO, HCl, HF, and SO2, at altitudes between 65 and 130 km has been demonstrated. We report results from a series of six occultations with observations of the 4-mm SO2 band at latitudes 69°–88°N and 23°–30°N. It is the first time when the vertical distribution of SO2 is retrieved above the clouds with the help of solar occultation direct method. The sulfur dioxide transmission spectrum is measured on a background of strong CO2 absorption. Each retrieved vertical profile of SO2 is characterized by few points; the mixing ratio of SO2 being 0.1 ppm with scale height 1 ± 0.4 km for polar measurements (evening observations) and 1 ppm with scale height 3 ± 1 km at low latitudes (morning observations) at the altitude of about 70 km. Upper limits of <0.05 ppm are established around 75 km. Citation: Belyaev, D., O. Korablev, A. Fedorova, J.-L. Bertaux, A.-C. Vandaele, F. Montmessin, A. Mahieux, V. Wilquet, and R. Drummond (2008), First observations of SO2 above Venus’ clouds by means of Solar Occultation in the Infrared, J. Geophys. Res., 113, E00B25, doi:10.1029/2008JE003143.
1. Introduction [2] Sulfur dioxide is an important component of Venus’ atmosphere because this gas is the chemical precursor of H2SO4 droplets clouds which completely enshroud Venus. Any significant change in SO2 above and within the clouds of Venus can affect photochemistry and dynamics of the clouds [Esposito et al., 1997]. Moreover, sulfur dioxide behavior in the atmosphere may be an indicator of possible geological activity on the surface of the planet. [3] Sulfur dioxide in the Venusian atmosphere was first detected from ultraviolet Earth-based observations of the solar light scattered by cloud particles and Rayleigh scattering with a high spectral resolution 0.1 nm [Barker, 1979], and the detected mixing ratio was 0.02 – 0.5 ppm at the cloud top. Space-based identifications of SO2 UV absorption followed soon from Pioneer Venus orbiter (PV) [Stewart et al., 1979] and International Ultraviolet Explorer (IUE) [Conway et al., 1979]. The first one provided a retrieval of SO2 vertical column density of about 1017 cm2 with a 1
Space Research Institute, Moscow, Russia. Service d’Ae´ronomie du CNRS, Verrie`res-le-Buisson, France. 3 Also at Institut Pierre Simon Laplace, Universite´ de Versailles-SaintQuentin, Guyancourt, France. 4 Belgian Institute for Space Aeronomy, Brussels, Belgium. 2
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003143$09.00
spectral resolution of 1.3 nm [Esposito et al., 1979] while the second one registered a column density of 4 1016 cm2 (0.1 – 0.8 ppm above the clouds) with a higher resolution of 0.3 nm. These abundances were larger by orders of magnitude than previously established upper limits delivered from rocket soundings in 1967 [Jenkins et al., 1969; Anderson et al., 1969] and Orbiting Astronomical Observatory (OAO) in 1972 [Owen and Sagan, 1972]. Continuous observations of the PV UV spectrometer from 1978 to 1986 showed a steady decline of the cloud top SO2 content down to 20– 50 ppb [Esposito et al., 1988]. This decline was confirmed by 1987 – 1988 IUE observations [Na et al., 1990]. Inbetween observations of sulfur dioxide in the UV range were made in 1983 by a balloon-borne spectrophotometer with a resolution of 0.4 nm [Parisot et al., 1986]. These measurements gave a mixing ratio of 70 ppb and, hence, confirmed the decrease of SO2 at that moment. Later rocket measurements in the UV range with a resolution of 0.2 nm resulted in SO2 content of 80 ± 40 ppb in 1988 and 120 ± 60 ppb at the equator in 1991 [McClintock et al., 1994; Na et al., 1994]. That was in agreement with PV and IUE results. One more UV measurement was made in 1995 from Hubble Space Telescope (HST) [Na and Esposito, 1995]. The SO2 abundance was 20 ± 10 ppb that is lower than values obtained in 1978– 1991 by a factor of 2 – 5 again confirming the decrease of SO2. [4] Sulfur dioxide was also measured in the infrared spectral range by the Fourier Spectrometer on board Venera
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15 orbiter [Moroz et al., 1985]. The instrument detected SO2 absorption bands at 519, 1150, and 1360 cm1 with a spectral resolution of 5 – 7 cm1. Nadir measurements covered the high northern latitudes almost completely; the sounding at altitude was around 69 km (40 mbar pressure). A strong dependence of SO2 content on latitude was reported: from 20 ppb at equator to 0.5 ppm in polar region. These data were found consistent with Pioneer Venus and IUE results [Zasova et al., 1993], taking into account the different effective sounding altitudes. [5] The goal of this paper is to present the first results concerning the detection of sulfur dioxide above the Venus’ cloud top by means of solar occultation in the infrared range using SOIR on board Venus Express orbiter (VEX) [Titov et al., 2006; Bertaux et al., 2007] The use of solar occultations probing (at the planet terminator) allows us to measure the atmospheric transmission and to retrieve information about the vertical structure and the composition at altitudes between 65 and 130 km (see Vandaele et al. [2008] and Fedorova et al. [2008] for the description of the technique and results for other key constituents of the Venus atmosphere). The slant density of SO2 is rapidly decreasing with altitude because of photolysis by sunlight. In occultation geometry this gas originating from the clouds is, therefore, expected to be observed only near the cloud top. The high photometric accuracy of SOIR instrument allows to measure transmitted spectrum down to slant optical thicknesses of t 6 (from aerosols), making it possible to detect SO2 with a satisfactory signal/noise ratio down to 68 km.
2. Instrument [6] SOIR instrument is a compact high-resolution infrared echelle spectrometer, in which an acoustooptical tunable filter (AOTF) is used for preliminary sorting of echelle diffraction orders [Korablev et al., 2004; Nevejans et al., 2006]. The instrument with spectral range 2300– 4500 cm1 and resolution of 0.15 cm1 is capable to detect important minor gaseous constituents such as CO, SO2, HCl, HF, H2O and HDO by means of solar occultation. The detailed description of the spectrometer, its in-flight performance, and calibrations are given in papers by Nevejans et al. [2006] and Mahieux et al. [2008]. We describe briefly the main parameters of the instrument, and some details related to SO2 detection. [7] The spectrometer’s rectangular field of view (FOV) is 30 arc minutes in spatial dimension, by 2 arc min in spectral dimension. This FOV is projected into 256 by 320 elements of an HgCdTe matrix detector. The elements of the detector, corresponding to the height of the slit are binned to form two groups each corresponding to 15 arc minutes. The angular diameter of the Sun from Venus orbit is 40 arc min. The width of the slit corresponds to 2 arc min. The spectrum is dispersed by the echelle onto the 320 spectral elements of the detector. The resolution of the echelle spectrometer, operating in diffraction orders 104– 204 is 0.15 cm1. The SO2 absorption bands were studied in the 4-mm range (2500 cm1), corresponding to 111 and 112 diffraction orders. The AOTF serves as an electrically tuned filter with a passband of 25 cm1, which is slightly larger than the widths of a single diffraction order [Mahieux et al., 2008]. Moreover, the spectral profile of the AOTF trans-
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mittance with side lobes includes also a few adjacent orders. Therefore, the measured spectrum contains a mixture of several orders. In the present study we consider the contribution of five nearby diffraction orders from each side. Calibration of the AOTF spectral function is made as described by Mahieux et al. [2008].
3. Calculations 3.1. Modeling of Transmission Spectra [8] The solar occultation technique allows measuring the atmospheric transmission at the terminator of a planet (either sunrise or sunset). For a spacecraft moving along its orbit the Sun is occulted by the planetary limb. During the occultation the FOV of the spectrometer is maintained by the inertial spacecraft attitude in Sun direction. A solar spectrum, measured out of the atmosphere (in our case, above 150 km of tangent altitude), is considered as a reference spectrum F0(n) (n denotes wave number). Then, the instrument registers a set of solar spectra F(n, z), passed through the atmosphere at various altitudes of the target point z. The solar occultation method is self-calibrated with respect to the atmospheric transmission which for a given altitude z is determined as: T ðn; zÞ ¼ F ðn; zÞ=F0 ðn Þ:
ð1Þ
The strongest absorption band of SO2 within the spectral range of SOIR is located around 4 mm (2500 cm1). This range is contaminated by a strong absorption of O16C12O18 isotope of CO2 (20003 – 00001 transition) and the SO2 spectrum is registered on the background of the CO2 absorption (Figure 1). The contribution of other atmospheric gases such as H2O and its isotopes, HCl, HF is negligible in this range; among the molecules listed in HITRAN 2004 spectral database [Rothman et al., 2005] the strongest absorption is expected from HDO, and it does not exceed 102. The transmission may be then theoretically expressed by the following combination: T ðn; zÞ ¼ TCO2 ðn; zÞTSO2 ðn; zÞT t ðn; zÞ;
ð2Þ
where 2 TCO2 ðn; zÞ ¼ exp42
Z1
3 sCO2 ðT ðl Þ; pðlÞÞnCO2 ðl Þdl5;
z
2 TSO2 ðn; zÞ ¼ exp42
Z1
3 sSO2 ðT ðl Þ; pðl ÞÞnSO2 ðl Þdl5;
z
Tt ðn; zÞ ¼ exp½t a ðn; zÞ :
Integration is performed along the optical trajectory l in the atmosphere. Gaseous local density n(l) is defined from respective slant density using the ‘‘onion peeling’’ method [Be´zard et al., 1987; Vandaele et al., 2008] assuming spherical symmetry of the atmosphere, and hydrostatic equilibrium. The absorption cross section si(T, p) can be
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Figure 1. Synthetic spectra of SO2 (for several mixing ratios, 0.05 (red), 0.1 (blue), and 1 ppm (black)) are presented at occultation altitude of 70 km on a background of CO2 spectrum (green curve) in range 2470 –2525 cm1 corresponding to the echelle orders 111 and 112.
calculated using line-by-line simulations of gaseous absorption using line parameters from spectroscopic database. The slant optical thickness of aerosol t a(n, z) is determined from the occultation procedure as
contribution of adjacent diffraction orders of the echelle spectrometer. Separate CO2 and SO2 simulations with the orders mixing are shown as well (green and black curves, respectively).
t a ðn; zÞ ¼ lnðTa ðn; zÞÞ;
3.2. Algorithm of Fitting [11] As it is evident from the synthetic model (Figures 1 and 2), the SO2 absorption features are overlapped by a strong CO2 structure. Not a single isolated line of SO2 is available for the analysis. Therefore, a combined synthetic model TCO2TSO2 including extinction and the mixing of diffraction orders was compared to the measured spectrum at each altitude. The fitting was performed in the spectral range of 2483 –2500 cm1 in the diffraction order 111, and of 2505 – 2523 cm1 in the order 112. The procedure to compute the synthetic model including aerosol extinction and the order mixing consists of the following steps: [12] 1. Computing of theoretical TCO2(n, z) and TSO2(n, z) transmissions separately for a specified altitude z. The CO2 density nCO2 and the SO2 mixing ratio fSO2 = nSO2/nCO2 are considered as variable parameters. [13] 2. Combined transmission TCO2SO2(n, z) = TCO2(n, z) TSO2(n, z) is calculated taking into account the mixing of echelle orders with the contribution of five diffraction orders ord (n, z). The full from each side: TCO2SO2(n, z) ! TCO2SO2 spectral range which contributes to the combined spectrum amounts to 2370 – 2640 cm1. mod data (n, z) and Idif (n, z) [14] 3. Two ‘‘differential’’ continua Idif ord are extracted from both theoretical TCO2SO2(n, z) and measured Tdata(n, z) spectra (equation (1)), in a similar manner as in the method of differential optical absorption spectroscopy (DOAS) [Plat, 1994]. Continuum curves are determined at a number of spectral intervals between absorption lines where absorption is minimal. Square polynomial interpolation is used in between.
ð3Þ
where Ta(n, z) is the continuum derived from the measured transmission spectra (equation (1)) following a retrieval procedure described below. In the present analysis we do not take into account the atmospheric refraction, which may introduce an error of a fraction of km into the effective sounded attitude at 70 km [Mahieux et al., 2008]. [9] All necessary spectroscopic gaseous parameters for calculations in equation (2) are taken from HITRAN 2004 database. Within the spectral range of interest (2450 – 2550 cm1) the SO2 absorption is observed on the background of carbon dioxide absorption dominated by isotopic band of 16O12C18O (628), and the mixing ratio of SO2 is derived with respect to this isotope. The isotopic ratio of 628 isotope on Earth is 3.9 103, and it is almost the same on Venus with an accuracy of 5% (as discussed by Vandaele et al. [2008]). Further in the text, for simplicity, CO2 will signify the 628 isotope. [10] A set of SO2 synthetic spectra of Venus atmospheric transmission in occultation geometry for the tangent altitude of 70 km is presented together with the CO2 model (including all isotopes) in Figure 1. We assumed VIRA-2 empirical atmospheric model [Moroz and Zasova, 1997; Zasova et al., 2006] for atmospheric density and uniformly mixed SO2 with several relative abundances: 0.05, 0.1 and 1 ppm (from the VIRA data this value is 0.1 ppm for the same altitude). In Figure 2 the spectrum of the combined TCO2TSO2 transmission (blue curves for 111 and 112 orders) at the altitude 68 km is modeled taking into account
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Figure 2. Synthetic spectra of combined transmission (CO2*SO2) (dashed curves) are presented for echelle orders (top) 111 and (bottom) 112 at the altitude of 68 km together with separated models of CO2 (thin curves) and SO2 (for 1 ppm; thick curves). All the spectra are calculated taking into account the contribution from ±5 adjacent echelle orders due to nonzero side lobes of the AOTF passband function. [15] 4. Calculation of additional extinction factor (due to mod (n, z) that attenuation in the atmosphere) A(n, z) = 1/Idif describes difference between gaseous ‘‘differential’’ and aerosol continua. The aerosol continuum is then defined as data (n, z) A(n, z), and, therefore, we can find the Taer(n, z) = Idif atmospheric optical thickness t a(n, z) following equation (3). [16] 5. Finally, a combined synthetic transmission is ord (n, z)Tt(n, z) (equation (2)). found as Tmod(n, z) = TCO2SO2 [17] The resulted synthetic model Tmod(n, z) is fitted to each measured spectrum to retrieve the values of nCO2 and fSO2 corresponding to the minimum of the function S ðnCO2 ; fSO2 Þ ¼
X Tdata ðn i ; zÞ Tmod ðn i ; zÞ2 si
i
;
ð4Þ
where i is a number of a spectral point (1 – 320), si is the error corresponding to this point. The aerosol optical thickness t a is defined as a by-product of the retrieval at
each altitude. We will use value t a = 1 as reference to define the cloud top level (though it is rather the top of the haze, above the classical cloud deck).
4. Occultations 4.1. Strategy of Observations [18] From December 2006 to August 2007 more then 20 occultations covering the spectral range relevant to SO2 were performed. We have analyzed a subset of six occultations available before July 2007. Two measurements correspond to low latitudes (23°– 30°N) and four to high latitudes of Venus (69° – 88°N) (see Table 1). Observations nearby the pole were made from the closest distance to the limb (2000 km, that corresponds to altitude resolution about 1 km), while observations at low latitudes were made from larger distances (>5000 km with vertical resolution of >3 km). During an occultation two spectra are recorded simultaneously by two bins of the detector corresponding to
Table 1. Tabulation of Analyzed Occultations With Main Parameters of Observationa Orbit (in/eg) 235 244 251 262 434 435
(in) (in) (in) (in) (eg) (eg)
Date
Local Time
Lat
Lon
Dist, km
12 Dec 2006 21 Dec 2006 28 Dec 2006 8 Jan 2007 29 Jun 2007 30 Jun 2007
0149 1953 1916 1830 0558 0558
88° 85° 82° 69° 30° 23°
309° 246° 260° 282° 265° 267°
2130 2242 1741 2274 5502 6179
Spectral Range, cm1 (Order) 2483 – 2500 2483 – 2500 2505 – 2523 2483 – 2500 2483 – 2500 2505 – 2523
(111) (111) (112) (111) (111) (112)
Altitude Range, km
Altitude of t = 1, km
fSO2, ppm
HSO2, km
68 68 68 – 71 69 – 72 70 – 76 72 – 78
76 75 76 75 83 83
0.5 0.7 0.3 0.7 1 3
— — 1.1 0.8 2.8 3.3
a Orbit (in/eg), number of orbit (ingress/egress); Date, date of observation; Local time, local time on Venus; Lat, latitude; Lon, longitude; Dist, distance to limb; Spectral range (order), spectral range (cm1) with corresponded echelle order; Altitude range, altitude range of SO2 detection (km); Altitude at t = 1, altitude at which optical thickness is one; fSO2, mixing ratio of SO2 (ppm) at level of 69 km; HSO2, scale height of SO2 at 70 km.
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Figure 3. An example of transmission spectrum measured during an occultation on the orbit 244 at the altitude of 68 km is presented for the echelle order 111 (blue curve) with level of noise (red curve). ord (n, z) Tt (n, z) (see step 5 of the Simulated spectrum of combined transmission Tmod (n, z) = TCO2SO2 fitting algorithm) corresponds to the best fit of 0.7 ppm SO2 mixing ratio (black curve). Also, a fit in CO2 (n, z)Tt (n, z) is which SO2 absorption is forced to be zero and only the CO2 absorption remains Tmod shown by the green curve for the reference.
both parts of the spectrometer slit (see above). For the six occultations selected in present study the spectra measured in the two bins coincide within 10%. It means that for these occultations the slit is kept nearly parallel to the limb, and we have averaged spectra from the two bins. For the remaining observations the slit cannot be considered parallel to the limb; more sophisticated geometry calculations and appropriate data processing for this case will be addressed to a future paper about SOIR performance. 4.2. Data Analysis [19] Figure 3 displays a transmission spectrum measured at the altitude of 68 km during the occultation at orbit 244 (blue line). The spectrum is compared with a synthetic CO2 (n, z)Tt (n, z)) and of combined spectra of CO2 (Tmod CO2*SO2 (Tmod(n, z)) with retrieved optimal values of nCO2 and fSO2 as described above. The observed transmission is of the order of 0.3% only: the solar flux is highly attenuated, mainly by aerosols, with a horizontal optical thickness t a 5.8. The two synthetic spectra (for CO2 and CO2*SO2), which are built to mimic the observations, contain an aerosol attenuation, and the contribution of the adjacent orders, not completely eliminated by the AOTF function. This contribution amounts to 50% of the transmission in the main order observed. [20] The signature of SO2 in the measured spectrum may be extracted by dividing the observed transmission by the CO2 simulation using: SO2 Tdata ðn; zÞ ¼
Tdata ðn; zÞ CO2 ðn; zÞTt ðn; zÞ Tmod
ð5Þ
and comparing the result with what is obtained using the simulated transmission: SO2 Tmod ðn; zÞ ¼
Tmod ðn; zÞ CO2 ðn; zÞT ðn; zÞ Tmod t
ð6Þ
This is illustrated in Figure 4. The resulting model transmission of SO2 is slightly different from the one displayed on Figure 1, because here the contribution of adjacent orders leaking through the AOTF orders is taken into account. [21] An alternate estimation of SO2 contents was performed by comparing the measured SO2 transmission spectrum (equation 5) with the simulated one (equation 6) using SO2 a ‘‘scatter plot’’ where each point has for x coordinate Tmod SO2 (n, z), and for y coordinate the Tdata (n, z) (Figure 5a and equation (4)). The same exercise was performed with differential transmission dT, where dT = T – average(T(n)) (Figure 5b), where the average value is a sliding average over 1 cm1 along the whole 2484 – 2500 cm1 range. When the slope on the plot is one at some variable parameter fSO2, it means that the SO2 mixing ratio during the measurement is equal (or very near) to this fSO2. The slope is found using least squares method. In the example shown, the slope is k = 0.94 ± 0.20 for T, and k = 0.90 ± 0.16 for d(T). 4.3. Results [22] Vertical profiles of the sulfur dioxide mixing ratio are presented in Figure 6, together with aerosol slant optical thickness retrieved from the same measurement sequences as described above. Sometimes the fitting process does not allow finding a distinct minimum; in such cases only upper
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SO2 Figure 4. SO2 spectrum Tdata (n, z) (thin curve) retrieved from the same measured spectrum as in SO2 (n, z) (thick curve) following Figure 3 following equation (5) is compared with SO2 synthetic model Tmod equation (6). Arrows mark the most prominent absorption features of sulfur dioxide. Only the gaseous transmission without contribution of aerosol extinction is presented.
limits of the SO2 abundance can be found (indicated by triangles). All the remaining points signify a positive detection with error bars. For all four high-latitude occultations available (69° – 88°N) about 0.1 ppm of SO2 is detected at the level of 70 km, rapidly decreasing with a scale height of HSO2 = 1 ± 0.4 km above. At low latitudes (23° –30°N) we detect significantly higher contents of about 1– 3 ppm and HSO2 = 3 ± 1 km. We used only the detected SO2 values to estimate the scale height; upper limits were not taken into account. [23] The clouds top level seems to be a bit higher at low latitudes where mixing ratio of SO2 is found larger. The
altitude where t a = 1, a possible definition of clouds top level, is found to be of the order 75– 76 km for latitudes >69° and 83 km for latitudes 23° and 30°. Usually the clouds top is known to be at a lower altitude for high latitudes [Crisp, 1986]. It must be recognized that the clouds top is usually defined with a vertical depth of unity, at a much lower level than where the horizontal optical thickness is t a = 1. In fact, at this level we probe the upper haze, which stands above the main clouds. [24] In transmission we are sensitive to the total slant density of SO2 molecules, and not to the mixing ratio. This is why we would not be able to see SO2 at higher
Figure 5. ‘‘Scattered plots’’ illustrating the retrieval of SO2 transmission are presented for the same SO2 plotted versus the modeled one measurements as in Figures 3 and 4. (a) The measured transmission Tdata SO2 Tmod , with a linear regression of this dependence (solid line). (b) The transmission dT, where dT = T – average(T) is plotted in the same manner. When the slope of the linear approximation k is close to unity, the measured and theoretical SO2 contents are equal to each other (or very near). 6 of 10
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Figure 6. Vertical profiles of SO2 mixing ratio retrieved from all analyzed occultations (points with error bars) together with aerosol profiles (solid curves). Triangles denote the point where only the upper limits of the SO2 detection. The conditions of observations are indicted on the plots: d, distance to limb; Lat, latitude; Lon, longitude; LT, local time (see Table 1). altitudes, even if the SO2 mixing ratio would be >1 ppmv above 75 km.
5. Uncertainties [25] There are several aspects that impact on the accuracy of retrieved results. Some of them deal with specific features of the spectrometer and some others with uncertainties of the observation geometry. [26] The signal/noise ratio (SNR) of the instrument is about 103 for solar spectrum outside the atmosphere in the center of diffraction orders 111 – 112. On the edges of the detector SNR decreases by a factor of 4 – 5 due to combined effect of AOTF passband and echelle blaze functions. The extinction by aerosol and CO2 at the altitude of 68– 70 km (where SO2 is detected) decreases the signal by a factor of 5 103. The resulting SNR effective for the altitude range of SO2 detection varies from 2 to 3 in the center of the detector to 10– 15 at the edges. These uncertainties for each pixel are taken into account as si in equation (4), and they are translated into 40% of random error of retrieved SO2 abundance. [27] Systematic uncertainties considered are AOTF calibrations (different for the two parts of the slit), and altitude determination. The different AOTF calibrations [Mahieux et al., 2008] are minimized by averaging spectra from the two bins and can be estimated as 10% of relative error in the final gaseous content. The altitude uncertainties depend on the distance of the spacecraft to the limb (±0.5 km at
2000 km and ±2 km at 6000 km); also we have neglected the atmospheric refraction (see above). However, for the retrieval of the mixing ratio the altitude uncertainty results in extraction of CO2 local density with errors 5%. Summarizing all mentioned uncertainties we obtain errors of sulfur dioxide mixing ratio of about 50 –70%.
6. Discussions [28] The measurements of sulfur dioxide above clouds were performed during almost 40 years (Figure 7 and Table 2) with some gaps. Summarizing the history of SO2 observations on Venus clouds top, an immediate conclusion is that the SOIR measurements indicate higher mixing ratios by a factor 5 to 25 with respect to all previous measurements. The solar occultation measurements are made at the terminator of the planet; that is, around one half of the optical path is in the shade, and another is sunlit. Most of previous nadir and limb observations, with the exception Venera 15 IR measurements were performed in the UV range on the dayside of the planet. The increased contents of SO2 measured on the terminator may be explained by its photochemical behavior. Indeed, the photochemical lifetime of SO2 exposed to sun light at 68– 70 km altitude is 3 103 s (i.e., 1 h) [Mills et al., 2007], and its contents should depend on local time. However, other photochemical cycles in which the SO2 is involved are more complicated: with UV solar light SO2 is rapidly produced again from SO, and the timescale for net loss of SO2 via oxidation to SO3
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Figure 7. Measurements of SO2 on Venus above clouds, available from 1969 up to now (see Table 2). The SO2 content in ppb volume is considered at the level of 40 mbar (69 km of altitude). SOIR results indicate values for 68 and 70 km (only at high latitudes) and SPICAV UV results were derived from 100 to 110 km. is much longer, 106 – 107 s [Winick and Stewart, 1980; Krasnopolsky and Pollack, 1994; Pernice et al., 2004]. These estimations for the lifetime depends on the column abundance of SO2 and the production rate for SO3, both of which are not well constrained. [29] Comparing SOIR results on morning and evening terminators (see Table 1 and Figure 6), for morning occultations (at low latitudes) we measured 1 ppm of SO2 mixing ratio while in the evening (high latitudes) 0.1 ppm is detected at 70 km of altitude. Unfortunately, at present we do not have considerable statistics of observations to compare morning and evening occultations at similar latitudes. An extensive set of daytime and nighttime SO2 measurements was obtained by Fourier spectrometer (FS) on Venera 15. A strong latitude dependence of SO2 was retrieved from 20 ppb at the equator up to 0.5 ppm in polar regions [Zasova et al., 1993], but diurnal cycle dependencies were not analyzed. [30] Regarding the altitude distribution of SO2, on the average SOIR measurements indicate a scale height of 1 ±
0.4 km for polar measurements and 3 ± 1 km at low latitudes at the altitude about 70 km. The latter is comparable with the scale height retrieved by Venera 15 FS at 69 km [Zasova et al., 1993]. For polar regions the scatter of Venera 15 scale heights is large, data being clustered around 1 km and 4 – 6 km. The estimations of scale height from PV ultraviolet measurements [Esposito et al., 1984] ranges from 1.5 to 4 –5 km; the IUE measurements [Na et al., 1990] are consistent with a scale height of 3 ± 1 km. Our data are not in conflict with these previous measurements. The peculiarities of the solar occultation technique are operations at the planet terminator, hence, at local morning or evening, and a significant horizontal optical path, reaching 450 km at 70 km. Therefore, SOIR measurements might reflect evening/morning, and also horizontal inhomogeneities, the latter may be associated with small-scale haze structures. VIRTIS has observed such structures at a scale down to few kilometers [Piccioni, 2008]. [31] Several hypothesis have been put forward in order to explain the amazing evolution of sulfur dioxide content
Table 2. List of Previous Detections of Sulfur Dioxide Above Venus’ Clouds by Several Methods and in Several Spectral Rangesa Reference
Instrument
Spectral Range (nm)
SO2 Content (ppb)
Anderson et al. [1969] Owen and Sagan [1972] Esposito et al. [1988] Na et al. [1990] Parisot et al. [1986] Moroz et al. [1985]; Zasova et al. [1993] McClintock et al. [1994] Na and Esposito [1995] Bertaux et al. [2008] This paper
Rocket OAO UV spectrometer (PV) UV spectrometer (IUE) Balloon Fourier Spectrometer (Venera 15) Rocket HST SPICAV UV (VEX) SOIR (VEX)
200 – 300 200 – 360 200 – 320 200 – 220 200 – 320 7350, 8700, 19270
<50 <10 50 – 400 50 – 380 70 200
190 – 230 207 – 216 200 – 320 4000
80 – 120 20 300 – 1000 100 – 550
a
SO2 content correspond to altitude of 69 km except SPICAV UV results that were derived from 100 to 110 km.
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above Venus clouds on a timescale of tens of years from 1967 to 1995 (Figure 7). After a sharp increase between 1967 and 1979, a stable decline down to 20 ppb at 1995 was observed, with values close to upper limits of 1967. Three possible explanations of such behavior were recently recalled by Mills et al. [2007]: active volcanism [Esposito, 1984], changes in the effective eddy diffusion within the cloud layers [Krasnopolsky, 1986, p. 147], and changes in atmospheric dynamics [Clancy and Muhleman, 1991]. The volcano hypothesis uses the volcanic eruption as a source of buoyancy that allows the abundant SO2 below the Venus clouds to break through the stable upper cloud layer. The entrained SO2 is then observable remotely at the cloud top in the UV. Similarly, because the observed SO2 mixing ratio may differ by as much as 4 orders of magnitude from the base to the top of the cloud layers [Bertaux et al., 1996; Esposito et al., 1997], a small change in the effective eddy diffusion within the cloud layers may significantly alter the cloud top abundance of SO2 [Krasnopolsky, 1986, p. 147]. The average SO2 abundance below the clouds varies probably much more slowly, connected to the evolution of volcanic activity over geologic timescales [Fegley et al., 1997a]. [32] Large values of SO2 (up to 2 ppm) observed by SOIR on Venus Express in a limited set are supported by recent microwave observations by Sandor et al. [2007], indicating also a large variability of SO2, and by SPICAV UV stellar occultations with retrieval from 0.3 to 1 ppmv of the gas at 100– 110 km [Bertaux et al., 2008]. At present it would be premature to conclude that in 2004– 2007 we observe another global increase of SO2 contents, like one suspected in between 1970 and 1977 or a leftover of such an increase in the past. Sulfur dioxide content at Venus’ clouds top might express significant variations in the haze cover. Sulfuric acid droplets are a potential reservoir of SO2, and a surge of SO2 may results from an increased evaporation of these droplets, triggered by a temperature change of dynamical origin. [33] Further monitoring of SO2 by SOIR and SPICAV at high altitudes may clarify these dependences. [34] Acknowledgments. We thank our collaborators of the three institutes for the design and fabrication of the SOIR instrument, which was mainly built in Belgium by OIP Company, under the direction of IASB-BIRA (IASB-BIRA/Belgium, Service d’Ae´ronomie/France, and IKI/ Moscow). Russian team acknowledges RFBR grant 06-02-72563. Belgium team was supported by the Belgian Federal Science Policy Office and the European Space Agency (ESA, PRODEX program, contracts C 90268, 90113, and 17645). Procurement of AOTF was funded by CNES.
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Be´zard, B., J.-P. Baluteau, A. Marten, and N. Coron (1987), The 12C/13C and 16O/18O ratios in the atmosphere of Venus from high-resolution 10-mm spectroscopy, Icarus, 72(3), 623 – 634, doi:10.1016/0019-1035(87)90057-1. Clancy, R. T., and D. O. Muhleman (1991), Long-term (1979 – 1990) changes in the thermal dynamical and compositional structure of the Venus mesosphere as inferred from microwave spectral observations of 12 CO, 13CO and C18O, Icarus, 89, 129 – 146, doi:10.1016/0019-1035(91) 90093-9. Conway, R. R., R. P. Mccoy, C. A. Barth, and A. L. Lane (1979), IUE detection of sulfur dioxide in the atmosphere of Venus, Geophys. Res. Lett., 6(7), 629 – 631, doi:10.1029/GL006i007p00629. Crisp, D. (1986), Radiative forcing of the Venus mesosphere. I - Solar fluxes and heating rates, Icarus, 67, 484 – 514, doi:10.1016/00191035(86)90126-0. Esposito, L. W. (1984), Sulfur dioxide: Episodic injection shows evidence for active Venus volcanism, Science, 223, 1072 – 1074, doi:10.1126/ science.223.4640.1072. Esposito, L. W., J. R. Winick, and A. I. Stewart (1979), Sulfur dioxide in the Venus atmosphere: Distribution and implications, Geophys. Res. Lett., 6(7), 601 – 604, doi:10.1029/GL006i007p00601. Esposito, L. W., M. Copley, R. Eckert, L. Gates, A. I. F. Stewart, and H. Worden (1988), Sulfur dioxide at the Venus cloud tops, 1978 – 1986, J. Geophys. Res., 93, 5267 – 5276, doi:10.1029/JD093iD05p05267. Esposito, L. W., J.-L. Bertaux, V. Krasnopolsky, V. I. Moroz, and L. V. Zasova (1997), Chemistry of lower atmosphere and clouds, in Venus II, edited by S. W. Bougher, D. M. Hunten, and R. J. Phillips, pp. 415 – 458, Univ. of Ariz. Press, Tucson. Fedorova, A., et al. (2008), HDO and H2O vertical distributions and isotopic ratio in the Venus mesosphere by Solar Occultation at Infrared spectrometer on board Venus Express, J. Geophys. Res., doi:10.1029/ 2008JE003146, in press. Fegley Jr., B., G. Klingelhofer, K. Lodders, and T. Widemann (1997a), Geochemistry of surface-atmosphere interactions on Venus, in Venus II, edited by S. W. Bougher, D. M. Hunten, and R. J. Phillips, pp. 591 – 636, Univ. of Ariz. Press, Tucson. Jenkins, E. B., D. C. Morton, and A. V. Sweigart (1969), Rocket spectra of ˚ , Astrophys. J., 157, 913 – 924, Venus and Jupiter from 2000 to 3000 A doi:10.1086/150123. Korablev, O. I., J.-L. Bertaux, I. I. Vinogradov, Y. K. Kalinnikov, D. Nevejans, E. Neefs, T. Le Barbu, and G. Durry (2004), Compact high-resolution echelle-AOTF NIR spectrometer for atmospheric measurements, in Proceedings of the 5th ICSO, Eur. Space Agency Spec. Publ., ESA SP-554, 73 – 80. Krasnopolsky, V. A. (1986), Photochemistry of the Atmosphere of Mars and Venus, 147 pp., Springer, Berlin. Krasnopolsky, V. A., and J. B. Pollack (1994), H2O-H2SO4 system in Venus’ clouds and OCS, CO, and H2SO4 profiles in Venus’ troposphere, Icarus, 109, 58 – 78, doi:10.1006/icar.1994.1077. Mahieux, A., et al. (2008), In-flight performance and calibration of SPICAV SOIR on board Venus Express, Appl. Opt., 47(13), 2252 – 2265, doi:10.1364/AO.47.002252. McClintock, W. E., C. A. Barth, and R. A. Kohnert (1994), Sulfur dioxide in the atmosphere of Venus, Icarus, 112, 382 – 388, doi:10.1006/ icar.1994.1192. Mills, F. P., L. W. Esposito, and Y. L. Yung (2007), Atmospheric composition, chemistry, and clouds, in Exploring Venus as a Terrestrial Planet, Geophys. Monogr. Ser., vol. 176, edited by L. W. Esposito, E. R. Stofan, T. E. Cravens, pp. 73 – 100, AGU, Washington, D. C Moroz, V. I., and L. V. Zasova (1997), VIRA-2: A review of inputs for updating the Venus International Reference Atmosphere, Adv. Space Res., 19(8), 1191 – 1201, doi:10.1016/S0273-1177(97)00270-6. Moroz, V. I., et al. (1985), Venera 15 and Venera 16 infrared experiment. 4. Preliminary results of spectral analyses in the region of H2O and SO2 absorption bands, Cosmic Res., Engl. Transl., 23(2), 202 – 211. Na, C. Y., and L. W. Esposito (1995), UV observations of Venus with HST (abstract), Bull. Am. Astron. Soc., 27, 1071. Na, C. Y., L. W. Esposito, and T. E. Skinner (1990), International Ultraviolet Explorer observation of Venus SO2 and SO, J. Geophys. Res., 95(D6), 7485 – 7491, doi:10.1029/JD095iD06p07485. Na, C. Y., L. W. Esposito, W. E. McClintock, and C. A. Barth (1994), Sulfur dioxide in the atmosphere of Venus, II. Modeling results, Icarus, 112, 389 – 395, doi:10.1006/icar.1994.1193. Nevejans, D., et al. (2006), Compact high-resolution spaceborne echelle grating spectrometer with acousto-optical tunable filter based order sorting for the infrared domain from 2.2 to 4.3 mm, Appl. Opt., 45(21), 5191 – 5206, doi:10.1364/AO.45.005191. Owen, T., and C. Sagan (1972), Minor constituents in planetary atmospheres: Ultraviolet spectroscopy from the Orbiting Astronomical Observatory, Icarus, 16, 557 – 568, doi:10.1016/0019-1035(72)90102-9.
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Parisot, J. P., P. Rigaud, and D. Huguenin (1986), Balloon observations of Venus from 200 to 320 nm, Astron. Astrophys., 166, 333 – 336. Pernice, H., et al. (2004), Laboratory evidence for a key intermediate in the Venus atmosphere: Peroxychloroformyl radical, Proc. Natl. Acad. Sci. U. S. A., 101, 14,007 – 14,010, doi:10.1073/pnas.0405501101. Piccioni, G. (2008), Observations of Venus by VIRTIS aboard Venus Express, paper presented at 37th COSPAR Scientific Assembly, Montre´al, Canada, 13 – 20 July. Plat, U. (1994), Differential optical absorption spectroscopy, in Air Monitoring by Spectroscopic Techniques, Chem. Anal. Ser., vol. 127, edited by M. W. Sigrist, pp. 27 – 84, John Wiley, New York. Rothman, L. S., et al. (2005), The HITRAN 2004 molecular spectroscopic database, J. Quant. Spectrosc. Radiat. Transfer, 96(2), 139 – 204, doi:10.1016/j.jqsrt.2004.10.008. Sandor, B. J., T. Clancy, G. H. Moriarty-Schieven (2007), SO and SO2 in the Venus mesosphere: Observations of extreme and rapid variation, paper presented at 3th DPS Meeting, Am. Astron. Soc., Orlando, Fla. Stewart, A. I., D. E. Anderson, L. W. Esposito, and C. A. Barth (1979), Ultraviolet spectroscopy of Venus: Initial results from the Pioneer Venus orbiter, Science, 203, 777 – 779, doi:10.1126/science.203.4382.777. Titov, D. V., et al. (2006), Venus Express: Scientific goals, instrumentation, and scenario of the mission, Cosmic Res., Engl. Transl., 44(4), 334 – 348, doi:10.1134/S0010952506040071.
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Vandaele, A. C., et al. (2008), Composition of the Venus mesosphere measured by Solar Occultation at Infrared on board Venus Express, J. Geophys. Res., doi:10.1029/2008JE003140, in press. Winick, J. R., and A. I. Stewart (1980), Photochemistry of SO2 in Venus’ upper cloud layers, J. Geophys. Res., 85, 7849 – 7860, doi:10.1029/ JA085iA13p07849. Yung, Y. L., and W. B. Demore (1982), Photochemistry of the stratosphere of Venus—Implications for atmospheric evolution, Icarus, 51(2), 199 – 247, doi:10.1016/0019-1035(82)90080-X. Zasova, L. V., V. I. Moroz, L. W. Esposito, and C. Y. Na (1993), SO2 in the middle atmosphere of Venus: IR measurements from Venera-15 and comparison to UV data, Icarus, 105(1), 92 – 109, doi:10.1006/icar.1993.1113. Zasova, L. V., et al. (2006), Structure of Venusian atmosphere from surface up to 100 km, Cosmic Res., Engl. Transl., 44(4), 364 – 383, doi:10.1134/ S0010952506040095.
D. Belyaev, A. Fedorova, and O. Korablev, IKI, 117997, 84/32 Profsoyuznaya Street, Moscow, Russia. (
[email protected]) J.-L. Bertaux and F. Montmessin, Service d’Ae´ronomie du CNRS, BP 3, F-91371 Verrie`res-le-Buisson, France. R. Drummond, A. Mahieux, A.-C. Vandaele, and V. Wilquet, Belgian Institute for Space Aeronomy, 3 avenue Circulaire, B-1180 Brussels, Belgium.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114, E00B34, doi:10.1029/2008JE003094, 2009
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Evidence for carbonyl sulfide (OCS) conversion to CO in the lower atmosphere of Venus Yuk L. Yung,1 M. C. Liang,2,3 X. Jiang,4 R. L. Shia,1 C. Lee,1 B. Be´zard,5 and E. Marcq5 Received 31 January 2008; revised 30 December 2008; accepted 9 January 2009; published 26 March 2009.
[1] The chemical regimes in the atmosphere of Venus vary from photochemistry in the
middle atmosphere to thermal equilibrium chemistry in the lower atmosphere and the surface. Many chemical cycles have been proposed, but few details about these cycles are fully verified by comparison between observations and modeling. Recent high-quality data of carbonyl sulfide (OCS) and CO from ground-based and Venus Express observations provide a unique opportunity to test our understanding of chemistry and transport in the lower atmosphere of Venus. The spatial distributions of OCS and CO in the atmosphere reflect a sensitive balance between chemistry and transport. On the basis of our updated photochemical model and winds from Lee et al.’s (2007) general circulation model, we study the chemistry and transport in a simplified two-dimensional chemistry-transport model. OCS is produced by heterogeneous reactions on the surface; the middle atmosphere is a net sink for OCS. The combination of data and modeling provides strong evidence for the loss of OCS by conversion to CO. The detailed chemical mechanism is currently unknown, although a number of speculations have been proposed. The sensitivity of the distributions of OCS and CO to model parameters is reported. Citation: Yung, Y. L., M. C. Liang, X. Jiang, R. L. Shia, C. Lee, B. Be´zard, and E. Marcq (2009), Evidence for carbonyl sulfide (OCS) conversion to CO in the lower atmosphere of Venus, J. Geophys. Res., 114, E00B34, doi:10.1029/2008JE003094.
1. Introduction [2] Sulfur chemistry is critical to the composition of the Venus atmosphere, and four sulfur species have been firmly identified: SO2, SO, OCS, and H2SO4 (vapor and aerosol). Mills et al. [2007] have recently carried out an extensive review of chemistry in the atmosphere of Venus, and the reader is referred to this paper for a summary of previous results. As noted in this review, there are two parts to the chemistry of sulfur species in the atmosphere of Venus. On the surface of Venus (and possibly in the dense, hot lower atmosphere) the chemistry is dominated by thermodynamic equilibrium chemistry [Fegley et al., 1997]. Near and above the cloud tops, the chemistry is driven by solar UV radiation. Thus, the partitioning of sulfur among the different species represents a competition between thermodynamic equilibrium chemistry at the surface and in the lower atmosphere and photochemistry in the middle atmosphere. This underscores the importance of transport and mixing in
1 Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, California, USA. 2 Research Center for Environmental Changes, Academia Sinica, Taipei, Taiwan. 3 Graduate Institute of Astronomy, National Central University, Jhongli, Taiwan. 4 Department of Earth and Atmospheric Sciences, University of Houston, Houston, Texas, USA. 5 LESIA, Observatoire de Paris, Meudon, France.
determining the distribution of chemical species in the atmosphere of Venus. [3] Early works on sulfur chemistry on Venus by Prinn [1975, 1978, 1979] is based on a prediction by Lewis [1970] of sulfur species with mixing ratios of 60 ppmv for OCS, 6 ppmv for H2S, and 0.3 ppmv for SO2. Prinn [1975] suggested a scheme of photochemical formation of sulfuric acid and polysulfur from carbonyl sulfide OCS. [4] The primary sulfur carrier from the deep atmosphere and surface to the middle atmosphere is OCS. Near and above the cloud tops, OCS readily dissociates, releasing the S atom. OCS þ hn ! CO þ S 1D
ð1Þ
where S(1D) is the first electronically excited state of the S atom and is 1.145 eV above the ground S(3P) state. The most likely fate of S(1D) is quenching: S 1D þ CO2 ! CO2 þ S 3P
ð2Þ
The other branch of (2) forming CO + SO is endothermic. If there is a source of oxygen atoms (such as photodissociation of CO2), S undergoes oxidation to SO2, which subsequently forms SO3 and H2SO4 and condenses to form the cloud layers as follows. [5] The S atom gets oxidized to SO by reacting with O and O2
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003094$09.00
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S þ O þ M ! SO þ M
ð3Þ
S þ O2 ! SO þ O
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Table 1. Mixing Ratios of CO2 and Trace Gases at the Lower Boundary of 58 km Species
Mixing Ratio
Reference
CO2 H2 O CO HCl SO2 OCS
0.965 2.6 105 3.0 105 4.0 107 1.3 104 3.0 107
Yung and DeMore [1982] Mills et al. [2007], Bertaux et al. [2007] Yung and DeMore [1982] Yung and DeMore [1982] Mills et al. [2007], Be´zard et al. [1993] Krasnopolsky [2008], this work
Further oxidation to SO2 can proceed via the three-body reaction SO þ O þ M ! SO2 þ M
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emphasis on carbonyl sulfide (OCS) and carbon monoxide (CO). Section 2 describes the mechanistic aspects of this work, consisting of three key components: (1) an updated 1-D photochemical model above the cloud tops, (2) a 2-D transport model based on winds and diffusivities derived from a general circulation model (GCM), and (3) a simplified chemical model for OCS and CO for the entire atmosphere. The main modeling results for OCS and CO, along with comparisons with recent observations, and model sensitivity are discussed in section 3. Section 4 examines the implications of the model, especially on the nature of the unidentified sink reaction for OCS. More speculative results are discussed in Appendix A. The conclusions are stated in section 5.
ð5Þ
2. Photochemistry and Transport Catalytic oxidation by ClO is also possible SO þ ClO ! SO2 þ Cl
ð6Þ
The rate for reaction (6) has been measured in laboratory studies [Sander et al., 2006], and reaction (6) accounts for 10– 20% of the loss of SO at 66– 80 km altitude in recent photochemical models [Pernice et al., 2004; Mills and Allen, 2007]. Note that the net result is the oxidation of S to SO2, and eventually to H2SO4. [6] The above scenario dominates primarily above the cloud tops (located at 60 km above the surface). Below the cloud tops, oxygen is scarce. In the absence of an oxygen source, S reacts with other S bearing species to form polysulfur, Sx. S atoms can react to produce S2 or S2 may be produced via coupled S-Cl chemistry involving chlorosulfanes [Mills and Allen, 2007]. Production of S3 is possible through successive addition reactions. S3 is the chemical analog of ozone, known as thiozone. As the number of sulfur atoms increases, the polyatomic sulfur compounds tend to have lower-saturation vapor pressures. It is convenient to name all sulfur species beyond S3 ‘‘polysulfur’’ or Sx. The production of Sx is part of what has been termed the ‘‘slow atmospheric sulfur cycle’’ [von Zahn et al., 1983], which is completed by decomposition reactions in the lower atmosphere. In the UV region Sx absorbs strongly, and it may be the principal constituent of the unidentified UV absorber in the upper atmosphere of Venus [Toon et al., 1982]. Krasnopolsky and Pollack [1994] and Krasnopolsky [2007] presented detailed models of OCS and Sx chemistry. However, critical components of their chemical model are not based on known laboratory kinetics. [7] Therefore, the large picture is that OCS is produced at the surface, transported to the middle atmosphere, where it is destroyed by photolysis above the cloud tops. However, between 1 and 10 bars (30 to 50 km), there may be an additional sink, whose identity is currently unknown and which is a major focus of this paper. The compelling observational evidences are (1) vertical gradient of OCS between 30 and 50 km, (2) latitudinal gradient of OCS in this region, and (3) complementary latitudinal gradient in CO, suggesting OCS to CO conversion. [8] In this paper, we report a combined study of photochemistry and transport in a two-dimensional (2-D) chemistrytransport model (CTM) for the atmosphere of Venus, with
2.1. 1-D Photochemical Model [9] The UV radiation is averaged diurnally before photolytic calculations, and the latitude dependence is carefully taken into account (see Liang et al. [2005] and references contained therein for a detailed description of the model.) The 1-D model [Yung and DeMore, 1982; Mills et al., 2007] is from 58 to 112 km. The vertical eddy mixing coefficients equal those obtained previously [Yung and DeMore, 1982]. Additional UV attenuation caused by absorbers estimated by Crisp [1986] is included. The chemical scheme used in this paper is taken from Mills et al. [2007]. The volume mixing ratios of CO2, H2O, CO, HCl, SO2, and OCS at the lower boundary are given in Table 1. For the other species, free escape through the lower boundary is allowed. The upper boundary is closed for all species. [10] Figure 1 shows the mixing ratios of CO2 and the most abundant trace gases in the model. The abundance of CO increases with height because there is a source of CO from photolysis of CO2 and OCS higher up. CO is removed by catalytic chemistry involving Cl radicals derived from HCl photolysis, which results in a decline of HCl with altitude. Both SO2 and H2O are removed by the formation of H2SO4. As H2SO4 aerosols are removed from the model, they represent a net sink of SO2 and H2O, thereby explaining their rapid decrease with altitude. OCS is destroyed by photolysis above the cloud tops. The secondary peak around 85 km is caused by the reaction S þ ClCO ! OCS þ Cl
ð7Þ
with the S atom ultimately derived from SO2 þ hn ! SO þ O
ð8Þ
SO þ hn ! S þ O
ð9Þ
Figure 2 shows the mixing ratios of O, O2, and O3. The mesosphere is a net source of oxygen from CO2 photolysis. Oxygen is rapidly consumed by catalytic recombination with CO and oxidation of SO2 to SO3, followed by formation of H2SO4. The latter is removed from the model. [11] Figure 3 presents the mixing ratios of major hydrogen species. H2O is the major carrier of hydrogen to the
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Figure 1. Vertical profiles of mixing ratios of CO2, CO, HCl, SO2, and OCS predicted by the 1-D photochemical model. The boundary conditions are summarized in Table 1.
middle atmosphere above the cloud tops. Photolysis of H2O produces H2 and radicals H, OH, HO2, and H2O2 (odd hydrogen). The odd hydrogen species play a minor role in the photochemistry of Venus. Reactive chlorine bearing radicals are shown in Figure 4. They are derived from HCl photolysis, which also contributes to the production of odd hydrogen. The chlorine bearing species play a major role in the stability of the CO2 atmosphere [Yung and
DeMore, 1982]. There is a possible coupling to sulfur chemistry but the kinetics is highly uncertain [Mills and Allen, 2007]. [12] The possibility of Sx formation is illustrated in Figure 5. The principal sources of S are photolysis of OCS and SO. In the absence of O2, formation of higher polymers of sulfur is possible. The column production rates are summarized in Table 2. The Sn species are in photo-
Figure 2. Same as Figure 1 for oxygen-bearing molecules, O2, O, and O3. 3 of 16
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Figure 3. Same as Figure 1 for hydrogen-bearing molecules, H2O, H2, H, OH, HO2, and H2O2.
chemical equilibrium; that is, production and loss (primarily by photolysis) are equal, except for S8. The model stops at S8, which is removed from the atmosphere in the model. The possible role of this chemistry in the lower atmosphere is discussed in Appendix A. [13] The 1-D photochemical model with full chemistry provides the basis for constructing a simplified chemical model for OCS and CO for the 2-D chemistry-transport model. This will be taken up in section 2.3.
2.2. Transport Model [14] The general circulation model (GCM) used to calculate the stream function uses the dynamical core of the Hadley Centre general circulation model [Cullen, 1993] with a 5° horizontal grid staggered as an Arakawa B grid [Arakawa and Lamb, 1977] from pole to pole and a 31-level hybrid sigma pressure vertical coordinate from surface to about 100 km. Forcing and dissipation are provided by linear parameterizations of the radiative forcing (cooling to space) and boundary layer (Rayleigh friction). A qualita-
Figure 4. Same as Figure 1 for chlorine-bearing molecules, Cl, Cl2, ClO, ClCO, and ClCO3. 4 of 16
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Figure 5. Same as Figure 1 sulfur-bearing molecules, SO, S, S2, S3, S4, S5, S6, and S7. tively realistic radiative forcing is used to drive an equatorial superrotation and low wave number planetary waves in the middle atmosphere. No topography or diurnal cycle is used, and no active chemical processes are considered within the GCM. We briefly summarize below the essential parts of the model and make a comparison with other model results available in the literature. Full details of the GCM and this experiment are given by Lee et al. [2007]. [15] When a similar radiative forcing is used, the circulation produced by other GCMs [Yamamoto and Takahashi, 2003; Hollingsworth et al., 2007; Herrnstein and Dowling, 2007] is at least as strong as the circulation produced by Lee et al.’s [2007] GCM shown here, suggesting the circulation is not unreasonable for the radiative forcing used. When a radiative forcing with less heating in the lower atmosphere is used [e.g., Hollingsworth et al., 2007] the circulation in the lower atmosphere is much weaker, but the superrotation is also much weaker and does not compare well to the observed Venus atmosphere. However, the middle atmosphere circulation still resembles a Hadley-like overturning. [16] Observations by Venus Express [Limaye, 2007] and Pioneer Venus [Schubert, 1983] suggest that the circulation, at least at cloud level, is stronger than in the GCM shown here. However, there are few observations at depth, making the calculations in the deep atmosphere somewhat speculative. These results suggest that the ‘‘age of air’’ calculation shown here (see later discussion) should be more accurate in the middle atmosphere than in the lower atmosphere, but physical and observational constraints suggests even the lower atmosphere values presented should be within an order of magnitude of the actual values. [17] The meridional circulation is calculated from wind and temperature from GCM [Lee et al., 2007]. The data from the GCM are in 5° 5° latitude-longitude resolution. It has 31 vertical levels from 92 bar to 103 bar. We choose the same method as that of Jiang et al. [2004] to calculate
the transformed meridional circulation for the Venus. First, the three-dimensional (3-D) meridional mass flux, y P (l, 8, p), is determined by y P ðl; 8; pÞ ¼
2p a cos 8 g
Z
p
V ðl; 8; p0 Þdp0
ð10Þ
o
where a is the Venus radius, l, 8, and p are the longitude, latitude and pressure, V is the meridional velocity, and g is the gravitational acceleration rate. Then we interpolate the 3-D meridional mass flux to isentropic surfaces, using a mass-conserving linear interpolation scheme [Juckes et al., 1994]. The 2-D isentropic mass stream function, y q (8, q), is derived by zonal averaging of the 3-D isentropic meridional mass flux, y q (l, 8, q), along isentropes. Finally, we interpolate the 2-D isentropic mass stream function, y q (8, q), to pressure coordinates and scale by the density to produce the pressure surface stream function, y P (8, p), which is used to drive the 2-D CTM. The resulting pressure surface meridional circulation, y P (8, p), shown in Figure 6 (left), is like that for the Hadley cell in the Table 2. Column Production Rates of Sulfur Species Above the Cloud Topsa Species
Principal Reactions
Column Rate
S S2 S3 S4 S5 S6 S7 S8
SO + hn, S2 + O S + S + M, S + OCS S + S2 + M S2 + S2 + M S + S4 + M S + S5 + M, S3 + S3 + M S + S6 + M S4 + S4 + M, S2 + S6 + M
8.2 1011 2.4 1011 1.6 1010 2.7 1010 7.3 109 6.1 109 1.6 109 5.2 109
a Sn, sulfur species. Units are in molecules cm2 s1. The principal source of Sn is COS photolysis, whose column rate is 7.5 1010 molecules cm2 s1. At higher altitudes, SO photolysis provides an additional source of S atoms.
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Figure 6. (left) Mass stream function y and (right) horizontal eddy mixing coefficient Kyy calculated from Lee et al. [2007]. Units are 109 kg s1 for y and 109 cm2 s1 for Kyy. Minimum Kyy is set to be 1010 cm2 s1, a value that is chosen to be large enough that small-scale transport (resulting from the derivation of y) is reduced but small enough that the transport remains dominated by y. terrestrial atmosphere. There is upwelling in the tropics, followed by downwelling in the high latitudes. The total mass flux is 3000 109 kg s1, a value that should be compared to the terrestrial value of 200 109 kg s1 [see Peixoto and Oort, 1992, Figure 7.19]. Since the atmosphere of Venus is about 100 times more massive than the terrestrial atmosphere, this implies that the atmosphere turns over in about a decade, as the Earth’s atmosphere turns over in a year. This result is consistent with the age of air experiment described later. Note that the latitudinal extension of the circulation is much greater than that for the terrestrial atmosphere owing to the near absence of the Coriolis effect. [18] The horizontal mixing coefficients, Kyy, is calculated on the isentropic surfaces Jiang et al. [2004, equation (2)]. Then we interpolate it from the isentropic surfaces to the pressure surfaces for use in the 2-D CTM. Figure 6 (right) shows a latitude-altitude profile of Kyy. The values for Kyy are larger in the midlatitudes than the tropics owing to the presence of enhanced wave activities in the midlatitudes. Although the transport in the atmosphere is dominated by advection, the effect of horizontal diffusion is non trivial in smoothing out gradients in chemical tracers. [19] To visualize the effect of the circulation on transport of chemical species in the atmosphere, we compute the age of air, a quantity that has proved to be extremely useful in understanding the stratosphere-troposphere exchange in the terrestrial atmosphere [Hall and Waugh, 1997]. The age of air is obtained by following the trend of an inert tracer, whose abundance at the surface increases linearly with time. Figure 7 shows the age of air in the lower atmosphere of Venus derived from the circulation shown in Figure 6. It
takes about 10 (Earth) years for an air parcel released at the surface to reach the middle atmosphere at the cloud tops. Since the radiative forcing does not include a diurnal cycle, the GCM does not simulate the large temperature variations seen above 90 km [Bertaux et al., 2007] associated with the variation in radiative forcing. As a result the age of air is only approximate in the upper atmosphere. [20] There is a simple physical interpretation of the mass stream function [see Shia et al., 1989, equation (12)]: y P (8, p) is the integral from latitude = 90 to 8 for the vertical mass flux across the pressure = p level, or the integral from the surface to pressure = p for the horizontal mass flux across the latitude = 8. The convention is positive for upward and northward flows. For later reference, it is convenient to divide the atmosphere into nine boxes, as shown in Figure 8. Using the values of the stream function from Figure 6 (left) and the formulas by Shia et al. [1989], we derive the appropriate mass fluxes in and out of each box in units of 109 kg s1. The numbers have been rounded off, and so the values are approximate. Note that the divergence of fluxes from each box is zero owing to conservation of mass. As we will show later, the central box is crucial for our investigation of the OCS and CO budget. This simple diagram will allow us to make a qualitative estimate of the chemical production and loss rates of OCS and CO in this region of the atmosphere from observed gradients in the concentrations of the tracers. 2.3. 2-D Chemistry-Transport Model With Simplified Chemistry [21] The 2-D Caltech/JPL photochemical model has been employed to simulate the meridional distribution of hydro-
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Figure 7. Age of air (Earth year) derived from the circulation shown in Figure 6. carbons in the atmosphere of Jupiter [Liang et al., 2005] and the isotopic composition of ozone in the terrestrial atmosphere [Liang and Yung, 2007]. In principle, such 2-D models are similar to that used for Titan [Lebonnois et al., 2001] and have been useful in modeling the interaction of
chemistry and transport in planetary atmospheres. As noted earlier, we obtain our transport component of the model from a GCM. [22] We extend the base model [Yung and DeMore, 1982; Mills et al., 2007], which sets the lower boundary at the
Figure 8. Mass fluxes in units of 109 kg s1 across the boundaries of nine boxes in the atmosphere of Venus. The boundaries at the bottom, top, and sides are impermeable. The numbers are taken and rounded off from Figure 6. 7 of 16
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Figure 9. Model profiles of (top) OCS and (bottom) CO, in units of ppmv, using the model that includes reaction (R11). See text.
cloud tops, to the entire atmosphere from 0 to 112 km. Above the planetary boundary (2 km), the vertical eddy mixing coefficients are uniformly set at 100 cm2 s1 and we assume that the transport in this region is dominated by the advection. In the planetary boundary, the vertical eddy coefficients are set at 104 cm2 s1, a value high enough simulating for the mixings/transport of matters between the troposphere and the boundary, where the mixing ratios of OCS and CO are set to be 23 and 16.6 ppmv, respectively. Since we are primarily interested in the chemistry of OCS and CO the lower atmosphere, the chemistry in the atmosphere above the cloud tops only serves as the upper boundary and may be greatly simplified. Thus, the chemical model is reduced to about 10 key reactions involving production and destruction of OCS and CO. Similar simplification has been carried out for the chemistry of ozone in the Earth’s atmosphere [Liang et al., 2006]. In order to successfully simulate the distribution of OCS and CO in the lower atmosphere, we have to postulate a loss reaction for converting OCS to CO via reaction with X: OCS þ X ! CO þ Y
ð11Þ
If we assume that R11 is a bimolecular reaction, then the loss coefficient is given by L11 ¼ k11 ½X
ð12Þ
where k11 is the bimolecular rate coefficient [see Yung and DeMore, 1999, chapter 3]. By trial and error, we found that L11 108 s1 in the region around 30 km. We will discuss the chemical and physical mechanisms for this additional sink for OCS later. The identity of X in (12) remains elusive. Candidate molecules include SO3, (SO)2 and S.
3. Model Results 3.1. Modeling Results and Comparison With Observations [23] Figure 9 shows the latitude-altitude plots of OCS and CO in the region from 30 to 40 km. The vertical distribution of OCS is consistent with its source at the surface, transport to the middle atmosphere, where it is destroyed. The falloff of OCS with altitude is not as large as the observed value. This may be due to inadequacy in the transport model or
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Figure 10. Vertical profiles of OCS computed by the 2-D model: global mean (solid curves) and at 65°N (dashed curves). (left) Loss by photolysis of OCS alone. (right) Additional sink (R11) is included. See text.
loss rate. According to the schematic in Figure 8, the export of OCS poor air from the tropics and via the downward portion of the Hadley cell results in lower OCS mixing ratios at middle and high latitudes. The pattern of CO is similar to that for OCS, except that the gradients are now reversed. In this case, there is a source of CO from the mesosphere and an additional source from reaction (R11). Comparison between model predictions and observations will be discussed as follows. [24] Figure 10 (left) shows the vertical profiles of OCS from our model with loss only by photolysis. Although the model correctly simulates the falloff of OCS above the cloud tops, the scale height of the OCS density at 30 km is close to the atmospheric scale height of 10 km (i.e., small vertical gradient in mixing ratio), in contradiction to the observed scale height of about 2 km (see below). The reason is that the photolysis coefficient of OCS falls rapidly below the cloud tops. In order to produce the correct scale height of OCS at 30 km, we include an additional parameterized sink for OCS in the model. The results are shown in Figure 10 (right) and are in agreement with the observations. [25] The falloff of the OCS profile around 30 km is a well-documented observation [see Marcq et al., 2005, 2006, and references therein]. The slope is dlog½OCS =dlogP ¼ 5 1
ð13Þ
where P is pressure. The corresponding scale height at this altitude is about 2 km. Higher up in the atmosphere, above 60 km, the OCS concentrations are in fair agreement with recent measurements by Krasnopolsky [2008]. The poorer
agreement at 70 km may be due to the unrealistic circulation above the cloud tops. [26] The latitudinal distribution of OCS at 33 km from our model is presented in Figure 11 (top). The Venus Express VIRTIS (asterisks) and ground-based telescope IRTF (diamonds) are shown by special symbols. Dotted curve represents the cosine function of latitude. The model is in satisfactory agreement with the data. The reasons for the increase in the mixing ratio of OCS away from the tropics were discussed earlier. The error bars are 1-s standard deviation intervals, for both VIRTIS and IRTF retrievals. [27] We assume that the variation of OCS at the probed altitude (33 km) was caused by vertical translations of the reference profile given by Pollack et al. [1993]. A more detailed discussion of the considered set of OCS vertical profiles is available in the work by Marcq et al. [2005, 2008]. Orbits 98, 110, 111, 134, 136, 258, and 277 of the Venus Express spectra, have been used, mostly in science case 2 (off-pericenter observations), because of the need of long integration times. The spectra were provided by the high spectral resolution IR channel from the VIRTIS instrument (VIRTIS-H). The resolving power R in the used order of dispersion (order 5) is close to 1500 between 2.42 and 2.46 mm. The interpretation of these spectra is still in progress, so the retrievals in upcoming publications may differ significantly, although the main trends should persist at least qualitatively. On the other hand, the Earth-based retrievals have already been published by Marcq et al. [2005, 2006]. The used spectra were acquired on 13 August 2004 using the order 3 of the SpeX spectrometer (R 2000
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Figure 11. (top) Model OCS distribution at 33 km (solid curve). The Venus Express (asterisks) and ground-based telescope IRTF (diamonds) are shown by symbols. Dotted curve represents the cosine function of latitude. (bottom) Similar results for CO. from 1.92 mm to 2.48 mm, which includes the whole 2.3 mm spectral window) at the NASA IRTF. Only the retrievals from slit positions 1 and 2, as defined by Marcq et al. [2006, Figure 1], are used. [28] Figure 11 (bottom) shows the model latitudinal distribution of CO at 36 km. The latitudinal structure of CO is the complement of that of OCS in Figure 11 (top). This gives rise to the hypothesis that the chemical destruction of OCS leads to the production of CO, as suggested by reaction (R11). We will return to this issue later. 3.2. Sensitivity of Model to Input Parameters [29] The agreement between the model and data shown in Figure 11 is far from perfect. It is important to investigate the model sensitivity of these results to input parameters. The most important input to the model is the stream function that is responsible for advection in the model. We note that the model profile of OCS in Figure 11 is flatter than the data; this is true also for CO (see Figure 11). One way we can enhance the curvature of the chemical species is to fine-tune the stream function and Kyy and the chemical reaction rates. There are many ways to modify the parameters in the model. After many trials, we conclude that the latitudinal gradients in OCS and CO are sharpened by (1) narrowing the width and increasing the amplitude of the Hadley cell, (2) enhancing the unidentified chemistry R11, and (3) reducing horizontal diffusion.
[30] The modifications of the standard model are summarized in Table 3. Model A is same as the standard model, except that the peak of the stream function is modified by the following factors. We have nine latitude boxes evenly spaced between south pole (SP) and equator (EQ). Let the stream function at box i be multiplied by a factor fi, where i = 1 is the SP box at 85° and i = 9 is the EQ box at 5°. In model A, we set f6 = 2, allowing it to fall back linearly to 1 at SP and EQ. The changes in the northern hemisphere, for i = 10 to 18, are a mirror image of those in
Table 3. Sensitivity of Latitudinal Gradients in OCS and CO at 30 km to Model Parametersa Model
D[OCS] (ppmv)
D[CO] (ppmv)
Parameter Change
Standard A B C D E F G
1.65 1.83 1.77 2.15 2.24 2.63 2.80 1.22
1.51 1.84 1.88 1.57 1.78 2.50 3.00 1.08
y mid 2 y mid 3 Same as B, R11 3 Same as C, y mid 3 + y trop 2 Same as D, Kyy/4 Same as D, Kyy/10 Same as Standard, R11 = 0
a
D[OCS] = [OCS](0°) – [OCS](60°); D[CO] = [CO](0°) – [CO](60°). The atmosphere is divided into three vertical regions (see Figure 8). Stream functions in the middle and bottom regions of the atmosphere are y mid and y trop, respectively.
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Figure 12. Same as Figure 11 for model F (see Table 3 and text).
the southern hemisphere. The purpose of this change is to make the Hadley cell narrower, so that we can sharpen the latitudinal gradient of OCS and CO. A measure of the latitudinal gradient is Dc = c(0°) – c(60°), where c is the mixing ratio of OCS or CO. In model A, we have D[OCS] = 1.83 ppmv, as compared to the value of 1.65 ppmv in the standard model. The corresponding values for D[CO] are 1.84 and 1.51 ppmv. Therefore, the gradients of OCS and CO in model A are sharper than those in the standard model. [31] Further accentuation of the gradients is achieved if we set f6 = 3 in model B (see Table 3). In model C, we enhance the chemical reaction (R11) by a factor of 3. The chemistry change has a larger effect on OCS than CO. In model D, we enhance the tropical stream function by a factor of 2; that is f9 = 2. Horizontal diffusivity acts to remove the latitudinal gradient in the model. model E studies the effect of reducing Kyy by a factor of 3. This results in increases of the absolute values of Dc in the model. This factor is increased to 10 in model F, which represents the model with the greatest gradients studied so far. Figure 12 presents a comparison of data and model results for OCS and CO in model F. The agreement is somewhat better than those in Figure 11, but crucial data at higher latitudes are lacking. The sharp edges in the model profiles are produced by advection and they could not be smoothed out owing to the low Kyy in this model.
[32] Finally, in model G we test the relative importance of chemistry (R11) versus boundary conditions in determining Dc in the model. This run is same as the standard model, except that tropospheric chemistry R11 = 0. We note that in this case there remains a residual gradient due to downwelling of air at middle and high latitudes from the stratosphere (see Figure 8), where OCS (CO) is low (high). 3.3. Is the Sum of OCS and CO a Constant? [33] The above discussion strongly suggests the possibility that the latitudinal variation in OCS and CO seen in Figure 11 is driven by the conversion of OCS to CO. In this case the sum of [OCS] and [CO] should be a constant. Figure 13 shows the data and model for [OCS] + [CO], demonstrating that the sum is approximately a constant. Inspection of Table 3 suggests that D[OCS] + D[CO] is roughly zero, further verifying that the changes are the result of OCS to CO conversion.
4. Implications of the Model 4.1. Sulfur Budget [34] Our model which is consistent with the observations suggests that the global rate of destruction of OCS is 23,000 Tg-S a1 (teragram of sulfur atom equivalence per year, i.e., 1 Tg-S a1 = 3 1010 moles a1 = 2 1034 OCS molecules a1), a value that should be compared to the total volcanic source of 10 Tg-S a1 for the Earth [see,
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Figure 13. Comparison of model OCS + CO and data. Solid and dashed curves are for standard model and model F, respectively. Symbols are for two observations. e.g., Seinfeld and Pandis, 1998]. The mean destruction rate is equivalent to 3.0 1012 molecules cm2 s1, which is 43 times the corresponding rate of OCS destruction above the cloud tops (see Table 2). The middle atmosphere is a net sink for OCS. We argue that the large implied source of OCS is unlikely to be supplied by volcanic emission. For example, a large volcano like Pinatubo delivered about 10 Tg-S into the terrestrial stratosphere in 1991 [Bluth et al., 1992]. To produce fluxes consistent with our model would require 2300 Pinatubo-sized eruptions per year. There is no geochemical or geological evidence in support of enhanced volcanic activities of this magnitude on Venus. [35] It is more likely that OCS is produced by heterogeneous reactions on the surface from CO and polysulfur (Sx), or between CO and CO2 and surface minerals (e.g., pyrite), as discussed by Mills et al. [2007, section 3.2]. The reaction S2 + CO ! OCS + S is favored by Hong and Fegley [1997] and Krasnopolsky [2007]. The CO ultimately comes from CO2 photolysis above the cloud tops and destruction of OCS in the middle atmosphere. Its mixing ratio remains fairly high, 105 near the surface [see Mills et al., 2007, Figure 3]. It is a challenge for the laboratory chemists to confirm and quantify these surface reactions. [36] We shall argue that the OCS sink predicted by the model is a direct consequence of the circulation of the atmosphere and the observed vertical and horizontal gradients in OCS, independent of the details of the model. Referring to the middle box in Figure 8, the mass flux entering the box is 8in = 1630 109 kg s1. Thus, the S mass flux entering this box is 8in [OCS]in 32/44. The mass fluxes out of this box are the upward flux at 50 km (where there is almost no OCS; see Figure 10) and the two horizontal fluxes 8out = 410 + 390 = 800 109 kg s1 at
midlatitudes. Therefore, the S mass flux out of the center box is 8out [OCS]out 32/44. Putting in approximate values from the data presented in Figure 11, [OCS]in = 3.3 ppmv and [OCS]out = 2.5 ppmv, we have the next sink for OCS equal to 76, 000 Tg-S a1, a value that is within a factor of 3 of our earlier number. The overestimate is not surprising, as we attribute the cause of the latitudinal gradient entirely to loss by R11, whereas comparison between model G and the standard model suggests that part of the gradient is caused by the upper boundary via the downward fluxes at middle to high latitudes. 4.2. Previous Work on the Unidentified Reaction [37] Krasnopolsky and Pollack [1994] were the first to propose chemical destruction of OCS in the lower atmosphere of Venus. Krasnopolsky [2007] quantitatively modeled the following reactions for destroying OCS in the lower atmosphere: H2 SO4 ! H2 O þ SO3
ð14Þ
SO3 þ OCS ! CO þ ðSOÞ2
ð15Þ
ðSOÞ2 þOCS ! CO þ SO2 þ S2
ð16Þ
———————————————————— net H2 SO4 þ 2 OCS ! H2 O þ CO2 þ CO þ SO2 þ S2 ð17Þ
According to Krasnopolsky’s [2007] estimate, [SO3] 5 1013 cm3 and k15 5 1023 cm3 s1 at 30 km, resulting
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in L15 = k15[SO3] = 2.5 109 s1. This is of the same order of magnitude but somewhat smaller than L11 from (11). Similar estimate shows that the second Krasnopolsky reaction (16) yields an L16 = k16[(SO)2] that is larger than L15. The combination of L15 and L16 is of the same magnitude as L11. [38] A major weakness of this scheme is that reactions (15) and (16) are not based on laboratory studies. Laboratory experiments are urgently needed to verify or disprove these reactions. In view of these uncertainties associated with reactions (15) and (16), it is reasonable to propose alternatives. Three chemical schemes are discussed as follows. 4.3. Other Possibilities for the Unidentified Reaction [39] We will briefly describe three additional possible mechanisms for the destruction of OCS on Venus. The first two are based on heterogeneous reactions on or in particles. The third mechanism is the photosensitized dissociation of OCS via the photochemistry of Sx. [40] Chen et al. [2007] measured the rate of oxidation of OCS to CO2 on the surface of hematite. The uptake coefficient (g) lies in the range 107 to 1011. The loss rate of OCS in the atmosphere can be estimated using the expression L11 ¼ 1=4gvANa
ð18Þ
where v isffi the mean speed of the impacting molecule = pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 8kT =pm, A is the mean surface area of each hematite grain, and Na is number density of hematite grains. Since L11 108 s1 and v 104 cm s1, ANa must be 105 to 101 cm2 cm3. For comparison, ANa is 3 106 in the terrestrial stratosphere for dust after the eruption of the El Chichon volcano [Michelangeli et al., 1991]. Therefore, while not ruled out, this reaction is probably too slow to have an impact for the destruction of OCS in lower atmosphere. [41] Dalleska et al. [2000] measured the rate coefficient for the oxidation of OCS by H2O2 in sulfuric acid solution. While the abundance of sulfuric acid is high on Venus, the limiting chemical may be H2O2, which is produced in small abundance at 80 to 90 km (see Figure 3). Thus, unless we have evidence otherwise, we consider this mechanism unlikely. [42] The third proposal is the photosensitized dissociation of OCS via the photochemistry of Sx. The idea is based on with analogy with the photosensitized dissociation of H2O in the troposphere of the Earth: O3 þ hn ! O2 þ O 1D
ð19Þ
H2 O þ O 1D ! OH þ OH
ð20Þ
In this case, the dissociation of H2O could not occur directly because the short-wavelength UV photons required to dissociate H2O (l < 200 nm) are absorbed by O2 and O3 in the stratosphere. However, the long-wavelength UV photons (l > 310 nm) can penetrate to the troposphere and photolyze O3 to produce O(1D), which then reacts with H2O
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[see Sander et al., 2006; Yung and DeMore, 1999, chapter 10]. We argue that, by analogy with (19) and (20) on Earth, the photochemistry of Sx could lead to photosensitized dissociation of OCS using long-wavelength photons that can penetrate to the lower atmosphere of Venus. The details of the Sx scheme are referred to in Appendix A. [43] The key reaction in this case is the reaction (R7) in Appendix A, OCS + S ! CO + S2. This reaction was considered by Prinn [1975] and Krasnopolsky [2007], and was measured by Lu et al. [2006] in the temperature range 298– 985 K. This is adequate for our modeling. The crucial requirement is that the number density of S atoms at 30 km must be 105 cm3. Without a source of S atoms from the Sx photochemistry in Appendix A, Krasnopolsky [2007] predicts [S] 10 cm3, which is too low to have impact on the OCS distribution. 4.4. Testable Hypotheses [44] The nature of the unidentified reaction in the lower atmosphere of Venus can only be resolved with better atmospheric measurements and laboratory experiments. Therefore, the proposed chemical schemes constitute testable hypotheses that are useful for guiding future experiments. 4.4.1. Krasnopolsky’s Scheme [45] Reactions (15) and (16) must be demonstrated in the laboratory. The existence of SO3 follows from the thermal decomposition of H2SO4, but the existence of (SO)2 is less obvious. Both must be measured in the atmosphere of Venus. 4.4.2. Heterogeneous Reactions [46] The two heterogeneous reactions discussed in section 4.3 were studied at room temperature. It is conceivable that they could be faster at higher temperatures. Also, there may be other types of heterogeneous reactions not considered here. A thorough laboratory search is needed. 4.4.3. Photosensitized Dissociation [47] A comprehensive study of the UV radiation and Sx photochemistry in the lower atmosphere is needed to make a realistic assessment of the concentration of S atoms at 30 km. Measurements of sulfur speciation as well as UV irradiance in the lower atmosphere would be valuable.
5. Concluding Remarks [48] A simple two-dimensional chemistry and transport (2-D CTM) model is used to study the spatial distributions of OCS and CO in the lower atmosphere of Venus. The residual circulation and horizontal eddy diffusivities are derived from winds from Lee et al.’s [2007] general circulation model. The Hadley circulation provides rapid transport in the lower atmosphere of Venus. Mixing between the surface and the cloud tops occurs in as short as 10 years, which is significantly shorter than previous estimates of this time constant. The results for the latitude-altitude distributions of OCS and CO from our 2-D CTM are compared to recent observations (see Figure 11). A total of eight models were studied to test the sensitivity of model results to input parameters, as summarized in Table 3. High-latitude data for OCS and CO are urgently needed to distinguish between the current models, shown in Figure 11 (standard model) and Figure 12 (model F), or a more realistic
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Figure A1. Schematic diagram illustrating the Sn chemistry in the lower atmosphere.
CTM. Further work is also required on the GCM used in this work in order to improve the simulated circulation in our model so that we may better constrain the effects of the unknown reaction(s). [49] There is compelling evidence for an unidentified chemical reaction that destroys OCS in the lower atmosphere of Venus around 30 km. There is strong evidence that this unidentified reaction converts OCS to CO (but not to CO2), as both data and model suggest that the sum of OCS and CO is roughly a constant (see Figure 13 and section 3.3). A major implication of the model is that the total loss rate of OCS in the lower atmosphere is about 23,000 Tg-S a1. This is a robust result that does not depend on the details of the 2-D CTM. The value can be derived from heuristic arguments based on the strength of the Hadley circulation and observed chemical gradients (see section 4.1). [50] The nature of the unidentified reaction remains controversial. A survey of recent literature and this work reveal three testable hypotheses: (1) Krasnopolsky’s [2007] reactions of OCS with SO3 and (SO)2 (see section 4.2), (2) heterogeneous reactions on hermatite or sulfuric acid particles (see section 4.3), and (3) photosensitized dissociation of OCS driven by absorption of soft UV by Sx (see section 4.3 and Appendix A). New laboratory experiments and observations are urgently needed to prove or disprove the proposed hypotheses (see section 4.4). We note that if the last speculative chemistry were confirmed, there would important implications for the modeling of the evolution of the terrestrial atmosphere [Kasting et al., 1989; Farquhar et al., 2002].
Appendix A: Hypothesis of an Innovative Chemical Cycle [51] In order to account for the additional loss of OCS below the cloud tops (see Figure 10 (right)), we propose an innovative photosensitized dissociation of OCS driven by
photolysis of Sn using photons at long wavelengths that are able to penetrate through the clouds 2½S3 þ hn ! S2 þ S ðR3Þ S4 þ hn ! S3 þ S ðR5Þ S2 þ S þ M ! S3 þ M ðR8Þ S2 þ S2 þ M ! S4 þ M ðR9Þ 2½OCS þ S ! CO þ S2 ðR7Þ net 2 OCS ! 2 CO þ S2
ðA1Þ
The details of this innovative chemistry are summarized in Table A1 and shown in Figure A1. The reactions (Rn) refer to the reaction set listed in Table A1. Note that the scheme has the following characteristics: (1) photolysis at UV wavelengths releases S from OCS, (2) S reacts with OCS to form S2 and subsequent reactions produces Sn, (3) photolysis of Sn occurs at near UV and longer wavelengths, producing S, and (4) cycle repeats via characteristic 2. [52] The net result is summarized by (9). This scheme is known as photosensitized dissociation (for a discussion of this process see, e.g., Yung et al. [1984, section 2]). The rate coefficient for the key reaction (R7) has recently been measured in the temperature range of 298 – 985 K by Lu et al. [2006] and their expression is given in Table 1. At 500 K, k7 = 5.41 1014 cm3 s1 is much smaller than the value 5.30 1013 cm3 s1 used by Prinn [1975] but is close to the value 6.29 1014 cm3 s1 estimated by Krasnopolsky [2007]. We must emphasize the speculative nature of the proposed chemistry. There remain major gaps in our knowledge of the chemistry of Sn in Table A1 and Figure A1; new laboratory studies are needed to close the gaps. [53] In a crude model we have for Sn around 30 km the following concentrations, [S] = 1 105 cm3, [S2] = 2 108 cm3, [S3] = 7 109 cm3, [S4] = 7 109 cm3, [S8] = 4 1014 cm3. We do not go beyond S8, which has no atmospheric sink (and therefore a large concentration). It is
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Table A1. List of Reactions Related to the Chemistry of Sn as Shown in Figure A1 Reaction
Rate Coefficienta 5
(R1) (R2) (R3) (R4)
OCS þ h ! CO þ S S2 þ h ! S þ S S3 þ h ! S2 þ S S4 þ h ! S2 þ S2
J1 = 3.81 10 J2 = 9.93 103 J3 = 1.15 J4 = 1.07 101
(R5)
S4 þ h ! S3 þ S
J5 = 1.07 101
(R7) (R8)
OCS þ S ! CO þ S2 S2 þ S þ M ! S3 þ M
(R9)
S2 þ S2 þ M ! S4 þ M
(R10)
S4 þ S4 þ M ! S8 þ M
(R11) (R12) (R13)
S3 þ S ! S2 þ S2 S4 þ S ! S3 þ S2 CO þ S þ M ! OCS þ M
k7 = 6.63 1020T2.57e1180/T k8 (0) = 1.00 1030 k8 (1) = 3.00 1011 k9 (0) = 2.20 1029 k9 (1) = 1.00 1011 k10 (0) = 1.00 1030 k10 (1) = 3.00 1011 k11 = 3.00 1011 k11 = 3.00 1011 k13 (0) = 4.00 1036 e1940/T k13 (1) = 1.00 1015
Reference Molina et al. [1981] Moses et al. [2002, and references therein] Moses et al. [2002, and references therein] Moses et al. [2002, and references therein]; Branching ratio estimated Moses et al. [2002, and references therein]; Branching ratio estimated Lu et al. [2006] Mills [1998] Mills [1998] Mills [1998] Mills [1998] Mills [1998] Mills [1998]; Estimated on the basis of analogy with O + CO + M ! CO2 + M
The photolytic coefficients are given for a diurnally averaged model at 30°N at the top of the atmosphere (s1). Two-body and three-body rate coefficients are given in units of cm3 s1 and cm6 s1, respectively; k(0) and k(1) refer to rate coefficient in the low and high pressure limit, respectively. a
transported to the surface, where surface reactions will recombine the sulfur and CO to form OCS, thus completing the cycling of OCS. Note that S3 has been tentatively identified by UV data from Venera 14, and Sx is a leading candidate for the unidentified UV and blue absorber [see Mills et al., 2007, sections 3 and 6]. The amount quoted for S3 by Maiorov et al. [2005] is 0.03 to 0.1 ppb. Bertaux et al. [1986] reported 5 to 25 ppmv of S8 from 25 to 45 km from analysis of UV data obtained by Venera 11 and 12. Note that 10 ppmv of S8 at 30 km corresponds to [S8] 1 1015 cm3, which is of the same order magnitude as what is needed in the model. However, the authors concluded that such values might be 2 orders of magnitudes too high. One resolution of the conflict is that the UV absorption in the lower atmosphere of Venus may not be homogeneous, creating gaps through which UV may penetrate to much deeper levels than an average model would predict. [54] Acknowledgments. We thank K. Baines, P. Drossart, J. Moses, C. Parkinson, V. Natraj, and X. Zhang for helpful discussions, D. Crisp for providing UV absorber profiles, and K. F. Li, D. Yang, and X. Zhang for assistance in preparing the manuscript. Special thanks are due to W. B. DeMore for a critical discussion of sulfur chemistry and S. Lebonnois and three anonymous referees for raising fundamental issues the resolution of which led to a much better paper. This research was supported by NASA grant NNX07AI63G to the California Institute of Technology. M. Liang was supported by NSC 97-2628-M-001-001 grant to Academia Sinica.
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Lee, C., S. R. Lewis, and P. L. Read (2007), Superrotation in a Venus general circulation model, J. Geophys. Res., 112, E04S11, doi:10.1029/ 2006JE002874. Lewis, J. S. (1970), Venus: Atmospheric and lithospheric composition, Earth Planet. Sci. Lett., 10, 73 – 80, doi:10.1016/0012-821X(70)90066-X. Liang, M.-C., and Y. L. Yung (2007), Sources of the oxygen isotopic anomaly in atmospheric N 2 O, J. Geophys. Res., 112, D13307, doi:10.1029/2006JD007876. Liang, M. C., et al. (2005), Meridional transport in the stratosphere of Jupiter, Astrophys. J., 635, L177 – L180, doi:10.1086/499624. Liang, M. C., et al. (2006), Isotopic composition of oxygen of stratospheric ozone, J. Geophys. Res., 111, D02302, doi:10.1029/2005JD006342. Limaye, S. S. (2007), Venus atmospheric circulation: Known and unknown, J. Geophys. Res., 112, E04S09, doi:10.1029/2006JE002814. Lu, C. W., Y. J. Wu, Y. P. Lee, R. S. Zhu, and M. C. Lin (2006), Experimental and theoretical investigation of rate coefficients of the reaction S (P-3)+OCS in the temperature range of 298 – 985 K, J. Chem. Phys., 125(16), 164329, doi:10.1063/1.2357739. Maiorov, B. S., et al. (2005), A new analysis of the spectra obtained by the VENERA missions in the Venusian atmosphere: I. The analysis of the data received from the VENERA-11 probe at altitudes below 37 km in the 0.44 – 0.66 mm wavelength range, Sol. Syst. Res., 39, 267 – 282, doi:10.1007/s11208-005-0042-1. Marcq, E., B. Be´zard, T. Encrenaz, and M. Birlan (2005), Latitudinal variations of CO and OCS in the lower atmosphere of Venus from nearinfrared nightside spectro-imaging, Icarus, 179, 375 – 386, doi:10.1016/ j.icarus.2005.06.018. Marcq, E., T. Encrenaz, B. Be´zard, and M. Birlan (2006), Remote sensing of Venus’ lower atmosphere from ground-based IR spectroscopy: Latitudinal and vertical distribution of minor species, Planet. Space Sci., 54, 1360 – 1370, doi:10.1016/j.pss.2006.04.024. Marcq, E., B. Be´zard, P. Drossart, G. Piccioni, J. M. Reess, and F. Henry (2008), A latitudinal survey of CO, OCS, H2O, and SO2 in the lower atmosphere of Venus: Spectroscopic studies using VIRTIS-H, J. Geophys. Res., 113, E00B07, doi:10.1029/2008JE003074. Michelangeli, D. V., M. Allen, and Y. L. Yung (1991), Heterogeneous reactions with NaCl in the El Chichon volcanic aerosols, Geophys. Res. Lett., 18, 673 – 676, doi:10.1029/91GL00547. Mills, F. P. (1998), I. Observations and photochemical modeling of the Venus middle atmosphere. II. Thermal infrared spectroscopy of Europa and Callisto, Ph.D. thesis, Calif. Inst. of Technol., Pasadena, Calif. Mills, F. P., and M. Allen (2007), A review of selected issues concerning the chemistry in Venus’ middle atmosphere, Planet. Space Sci., 55, 1729 – 1740, doi:10.1016/j.pss.2007.01.012. Mills, F. P., L. W. Esposito, and Y. L. Yung (2007), Atmospheric composition, chemistry, and clouds, in Exploring Venus as a Terrestrial Planet, Geophys. Monogr. Ser., vol. 176, edited by L. W. Esposito, E. Stofan, and T. Cravens, pp. 73 – 100, AGU, Washington, D. C. Molina, L. T., J. J. Lamb, and M. J. Molina (1981), Temperature-dependent UV absorption cross-sections for carbonyl sulfide, Geophys. Res. Lett., 8, 1008 – 1011, doi:10.1029/GL008i009p01008. Moses, J. I., M. Y. Zolotov, and B. Fegley (2002), Photochemistry of a volcanically driven atmosphere on Io: Sulfur and oxygen species from a pele-type eruption, Icarus, 156, 76 – 106, doi:10.1006/icar.2001.6758. Peixoto, J. P., and A. H. Oort (1992), Physics of Climate, 520 pp., Am. Inst. of Phys., New York.
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Pernice, H., et al. (2004), Laboratory evidence for a key intermediate in the Venus atmosphere: Peroxychloroformyl radical, Proc. Natl. Acad. Sci. U. S. A., 101, 14,007 – 14,010, doi:10.1073/pnas.0405501101. Pollack, J. B., et al. (1993), Near-infrared light from Venus’ nightside: A spectroscopic analysis, Icarus, 103, 1 – 42, doi:10.1006/icar.1993.1055. Prinn, R. G. (1975), Venus: Chemical and dynamical processes in the stratosphere and mesosphere, J. Atmos. Sci., 32, 1237 – 1247, doi:10.1175/1520-0469(1975)032<1237:VCADPI>2.0.CO;2. Prinn, R. G. (1978), Venus: Chemistry of the lower atmosphere prior to the Pioneer Venus mission, Geophys. Res. Lett., 5, 973 – 976, doi:10.1029/ GL005i011p00973. Prinn, R. G. (1979), On the possible roles of gaseous sulfur and sulfanes in the atmosphere of Venus, Geophys. Res. Lett., 6, 807 – 810, doi:10.1029/ GL006i010p00807. Sander, S. P., et al. (2006), Chemical kinetics and photochemical data for use in stratospheric modeling evaluation number 15, JPL Publ. 06 – 2, 523 pp. Schubert, G. (1983), General circulation and the dynamical state of the Venus atmosphere, in Venus, edited by D. M. Hunten et al., pp. 681 – 765, The Univ. of Ariz. Press, Tuscon. Seinfeld, J. H., and S. N. Pandis (1998), Atmospheric Chemistry and Physics: From Air Pollution to Global Change, John Wiley, New York. Shia, R. L., Y. L. Yung, M. Allen, R. W. Zurek, and D. Crisp (1989), Sensitivity study of advection and diffusion coefficients in a twodimensional stratospheric model using excess C-14 data, J. Geophys. Res., 94, 18,467 – 18,484, doi:10.1029/JD094iD15p18467. Toon, O. B., R. P. Turco, and J. B. Pollack (1982), The ultraviolet absorber on Venus: Amorphous sulfur, Icarus, 51, 358 – 373, doi:10.1016/00191035(82)90089-6. von Zahn, U., S. Kumar, H. Niemann, and R. G. Prinn (1983), Composition of the Venus atmosphere, in Venus, edited by D. M. Hunten et al., pp. 299 – 430, Univ. of Ariz. Press, Tucson. Yamamoto, M., and M. Takahashi (2003), Superrotation and equatorial waves in a T21 Venus-like AGCM, Geophys. Res. Lett., 30(9), 1449, doi:10.1029/2003GL016924. Yung, Y. L., and W. B. DeMore (1982), Photochemistry of the stratosphere of Venus: Implications for atmospheric evolution, Icarus, 51, 199 – 247, doi:10.1016/0019-1035(82)90080-X. Yung, Y. L., and W. B. DeMore (1999), Photochemistry of Planetary Atmospheres, Oxford Univ. Press, New York. Yung, Y. L., M. Allen, and J. P. Pinto (1984), Photochemistry of the atmosphere of Titan: Comparison between model and observations, Astrophys. J. Suppl. Ser., 55(3), 465 – 506, doi:10.1086/190963.
B. Be´zard and E. Marcq, LESIA, Observatoire de Paris, Baˆtiment 18, pie`ce 111, F-92195 Meudon CEDEX, France. X. Jiang, Department of Earth and Atmospheric Sciences, University of Houston, Houston, TX 77004, USA. C. Lee, R. L. Shia, and Y. L. Yung, Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA. (
[email protected]) M. C. Liang, Research Center for Environmental Changes, Academia Sinica, 128 Academia Road, Section 2, Taipei 115, Taiwan.
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Modeling the distribution of H2O and HDO in the upper atmosphere of Venus Mao-Chang Liang1,2 and Yuk L. Yung3 Received 31 January 2008; revised 30 June 2008; accepted 2 December 2008; published 24 February 2009.
[1] The chemical and dynamical processes in the upper atmosphere of Venus are poorly
known. Recently obtained vertical profiles of trace species from the Venus Express mission, such as HCl, H2O, and HDO, provide new information to constrain these processes. Here, we simulate these profiles, using the model we have developed and described in a related paper by Yung et al. (2008), with special emphasis on the modeling of H2O and HDO. A new mechanism, the photo-induced isotopic fractionation effect (PHIFE) of H2O and HCl, is incorporated into our model. The observed enhancement of HDO could be attributed to (1) preferential destruction of H2O relative to HDO via PHIFE and (2) escape of hydrogen that enhances the abundance of D and hence its parent molecule HDO. Over a wide range of the sensitivity of the results to the changes of the two mechanisms, we find that the observed profiles of HDO and H2O profiles cannot be explained satisfactorily by current knowledge of chemical and dynamical processes in this region of the atmosphere. Several conjectures to tackle the problems are discussed. Citation: Liang, M.-C., and Y. L. Yung (2009), Modeling the distribution of H2O and HDO in the upper atmosphere of Venus, J. Geophys. Res., 114, E00B28, doi:10.1029/2008JE003095.
1. Introduction [2] Venus provides a window of opportunity in the solar system for studying the end-member of water evolution. Its highly enhanced D/H ratio, compared with the terrestrial value, suggests that about one Earth ocean might have been lost [e.g., Kasting and Pollack, 1983; Donahue, 1999, and references therein] via nonthermal escape of hydrogen (such as hydrodynamic escape, charge exchange, and collisionally induced escape). Hydrogen is produced by the photolysis of water, a process that is known to preferentially destroy the light isotopologue, resulting in the enrichment of the heavy parent molecule in the atmosphere [Cheng et al., 1999]. The process of enhancing the abundance of parent molecules is similar in nonthermal escape of hydrogen, and hence over the course of Venusian history, the D/H ratio is enriched as compared with the primordial value. A combination of photolytic and nonthermal escape processes fractionates the ratio in a way faster than each of them alone. In this paper, we investigate the D/H ratio affected by the photoinduced isotopic fractionation effect (PHIFE) of H2O/HDO and HCl/DCl in the present atmosphere of Venus. The reader is referred to Miller and Yung [2000] for a detailed explanation of PHIFE. The observed D/H in water in the upper atmosphere from Venus Express, along with other 1 Research Center for Environmental Changes, Academia Sinica, Taipei, Taiwan. 2 Graduate Institute of Astronomy, National Central University, Jhongli, Taiwan. 3 Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, California, USA.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003095$09.00
molecules (HCl, HF), provides additional constraints for understanding the relative importance of two dynamical processes: atmospheric circulation and escape. The former conserves the bulk D/H ratio in an air parcel and the latter enhances the ratio. As a result, the D/H ratio in water has less latitude dependence, compared with the case involving hydrogen escape. To evaluate the processes quantitatively in the upper atmosphere of Venus, one-dimensional (1-D) and two-dimensional (2-D) models are used. The data used for the study are described by Bertaux et al. [2007], Vandaele et al. [2008], and Fedorova et al. [2008].
2. Models [3] One-dimensional photochemical models are used to simulate the vertical profiles of H2O/HDO, HCl/DCl, H/D, H2/HD, OH/OD, HO2/DO2, CO2, CO, O2, O, O(1D), O3, Cl, ClO, Cl2, ClCO, and ClCO3 in the upper atmosphere (58 – 112 km) of Venus. The current model is based on Yung and DeMore [1982] and Mills [1998a, 1998b], and a subset of the chemistry is selected from Yung et al. [2008] to account for the UV attenuation in the upper part of the atmosphere. The selected hydrogen/deuterium chemistry is summarized in Table 1. The PHIFE of H2O/HDO and HCl/ DCl are taken from Cheng et al. [1999] and Bahou et al. [2001], respectively. The rest of the hydrogen and deuterium chemical reactions are assumed to be isotopically neutral. The transport and boundary conditions are primarily taken from Yung and DeMore [1982] and Mills [1998a, 1998b]. The CO2 mixing ratio is relatively uniform at 0.96 throughout the entire atmosphere. The fixed mixing ratios of 1, 0.1, 0.15, and 0.0075 ppmv are used for H2O, HDO, HCl, and DCl, respectively, at the lower boundary. The selected
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Figure 1. Modeled vertical profiles of H2O (black) and HDO (gray) in 1-D model with prescribed winds depicted by solid and dashed curves (see text). Data are from Bertaux et al. [2007]. D/H ratio used to better represent the Venus Express profiles (Figure 1) is a factor of 2 higher than the bulk ratio, which is 0.05 for [HDO]/[H2O]. The rest of the species are transported downward at the lower boundary at a velocity determined by dynamics; the velocities are 0.05 and 0.016 cm s1 for 1-D and 2-D models, respectively. The upper boundary is impermeable to all species. This is chosen to be our reference case. A model with hydrogen escape is also run. In this model, species other than atomic H and D have zero flux at the upper boundary. The escape fluxes of H and D are, respectively, 3.5 106 and 3.1 104 atoms cm2 s1 [Gurwell and Yung, 1993] or higher [Donahue, 1999]. Sensitivity of the model results with respect to the changes of transport and upper boundary conditions is studied using the 1-D model. The UV radiation is averaged diurnally before photolytic calculations, and the latitude dependence is carefully taken into account. (See Liang et al. [2005, and references therein] for a detailed description of the model.) The modeled latitude for 1-D models is set at 45°N, where we think it should better represent the condition at the polar region if the time constant of the large-scale meridional circulation is not small compared with the time of vertical transport and the lifetimes of H2O and HDO against photolysis. The major effect on the selection of the modeled latitude is the total photolysis rates (or J values) for molecules. The effect on isotopic ratios can be ignored under the current uncertainties of models and observations. We test the isotopic effect at a higher latitude at 80°N, and the largest difference of 5%, compared with that at 45°N, occurs above 110 km. [4] The 2-D version [Liang et al., 2005] of the Caltech/JPL photochemical mode is employed to simulate the meridonal distribution of H2O/HDO and HCl/DCl in the upper atmosphere of Venus. The current model and the adopted transport are described in a companion paper [Yung et al., 2008]. We solve the model at latitudes from pole to pole and altitudes from 56 to 112 km. The chemical species and reactions are taken from the 1-D model
described in the previous paragraph. The boundary conditions are the same as the 1-D case with hydrogen escape. This is selected to be our reference 2-D model. [5] The vertical profile of H2O provides insight into transport of the upper atmosphere of Venus. The SOIR data, taken near the north pole, show that H2O mixing ratio decreases with increasing altitude until 85 km and then increases above that [Bertaux et al., 2007]. Since there is no known source and sink of this magnitude (50%) above the cloud tops at 50 km and below 100 km (where photochemical processes become important), the profile has to be caused by transport. Several ad hoc advections have been tested, and 1-D models are used for such sensitivity study. One proposal that fits the data is having a downward advection (0.3 cm s1) between 75 and 85 km and upwelling (0.5 cm s1) above until 95 km where the transport becomes downwelling (0.5 cm s1). This qualitatively agrees with the fact that there is a temperature inversion layer at 100 km. General circulation models [e.g., Lee et al., 2007] predict, in general, that air ascends at low latitudes and descends at high latitudes (Hadley cell). The heating at 100 km results from the wind profile (see later discussion on the relation between wind and temperature). The resulting H2O and HDO profiles from this prescribed advection (dashed curves) are shown in Figure 1.
3. Results [6] The photolysis of H2O/HDO and HCl/DCl tends to enhance their isotopic composition d, which is defined by the deviation of the ratio of an isotopically substituted species and its normal molecule from that of the prescribed standard d ½ D=½ H =ð½ D=½ H Þ0 1;
ð1Þ
where [D] and [H] are the concentrations of [HDO] or [DCl] and [H2O] or [HCl], respectively. The subscript ‘‘0’’ refers
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Table 1. List of Hydrogen and Deuterium Chemical Reactionsa Reaction (R22) (R25) (R41) (R42) (R46) (R47) (R49) (R50) (R51) (R52) (R53) (R54) (R55) (R56) (R58) (R61) (R63) (R64) (R65) (R66) (R67) (R68) (R70) (R75) (R78) (R86) (R90) (R91) (R95) (R104) (R105) (R106) (R115) (R116) (R127) (R128) (R130) (R139) (R140) (R141) (R142) (R143) (R144) (R146) (R147) (R149) (R150) (R151) (R152) (R153) (R154) (R155) (R156) (R157) (R158) (R159) (R160) (R163) (R164) (R168) (R169) (R171) (R172) (R173) (R174) (R175) (R176) (R177) (R178) (R179) (R181) (R182) (R184)
H2O ! H + OH HCl ! H + Cl O(1D) + H2O ! 2OH O(1D) + H2 ! H + OH H + O2 + CO2 ! HO2 + CO2 H + O3 ! OH + O2 O + HO2 ! OH + O2 O + OH ! O2 + H OH + CO ! CO2 + H OH + H2 ! H2O + H OH + O3 ! HO2 + O2 HO2 + O3 ! OH + 2O2 H + HO2 ? 2OH H + HO2 ! H2O + O H + HO2 ! H2 + O2 OH + HO2 ! H2O + O2 2H + CO2 ! H2 + CO2 H + HCl ! H2 + Cl OH + HCl ! Cl + H2O O + HCl ! OH + Cl Cl + H2 ! HCl + H Cl + OH ! HCl + O Cl + HO2 ! HCl + O2 ClO + OH ! HO2 + Cl Cl + H + M ! HCl + M ClCO + H ! HCl + CO H + Cl2 ! HCl + Cl Cl + HO2 ! OH + ClO ClCO3 + H ! CO2 + Cl + OH O + H2 ! OH + H 2OH ! H2O + O O(1D) + HCl ! Cl + OH O(1D) + HCl ! O + HCl O(1D) + HCl ! H + ClO ClO + H2 ! HCl + OH O + H + M ! OH + M H + OH + CO2 ! H2O + CO2 HDO ! H + OD HDO ! D + OH DCl ! D + Cl O(1D) + HDO ! OH + OD O(1D) + HD ! H + OD O(1D) + HD ! D + OH D + O2 + CO2 ! DO2 + CO2 D + O3 ! OD + O2 O + DO2 ! OD + O2 O + OD ! O2 + D OD + CO ! CO2 + D OD + H2 ! HDO + H OH + HD ! HDO + H OH + HD ! H2O + D OD + O3 ! DO2 + O2 DO2 + O3 ! OD + 2O2 D + HO2 ! OH + OD H + DO2 ! OH + OD D + HO2 ! HDO + O H + DO2 ! HDO + O D + HO2 ! HD + O2 H + DO2 ! HD + O2 OD + HO2 ! HDO + O2 OH + DO2 ! HDO + O2 D + H + CO2 ! HD + CO2 D + HCl ! HD + Cl H + DCl ! HD + Cl OD + HCl ! Cl + HDO OH + DCl ! Cl + HDO O + DCl ! OD + Cl Cl + HD !DCl + H Cl + HD !HCl + D Cl + OD ! DCl + O Cl + DO2 ! DCl + O2 ClO + OD ! DO2 + Cl Cl + D + M ! DCl + M
Rate Coefficient 6
J22 = 2.5 10 J25 = 1.8 106 k41 = 2.2 1010 k42 = 1.1 1010 k46 = 2.0 1031 (T/300)1.6; k8 = 7.50 1011 k47 = 1.4 1010 e470/T k49 = 2.9 1011 e200/T k50 = 2.2 1011 e120/T k51 = 1.5 1013 k52 = 5.5 1012 e2000/T k53 = 1.6 1012 e940/T k54 = 1.1 1014 e500/T k55 = 7.3 1011 k56 = 1.6 1012 k58 = 6.4 1012 k61 = 4.7 1011 e250/T k63 = 5.0 1029 T1.3 k64 = 1.5 1011 e1750/T k65 = 2.6 1012 e350/T k66 = 1.0 1011 e3300/T k67 = 3.7 1011 e2300/T k68 = 1.2 e510/T k70 = 1.8 1011 e170/T k75 = 1.1 1011 e120/T k78 = 1.0 1032 k86 = 1.0 1011 k90 = 1.4 1010 e90/T k91 = 4.1 1011 e450/T k95 = 1.0 1011 k104 = 9.9 1032 T6.5 e1460/T k105 = 4.2 1012 e240/T k106 = 1.0 1010 k115 = 1.4 1011 k116 = 3.6 1011 k127 = 1.0 1012 e4800/T k128 = 1.3 1029 T1 k130 = 7.7 1026 T2
(1/2)J22
(1/2)J22
J25 k41 k42 k42 k46 k47 k49 k50 k51 k52 (1/2)k52 (1/2)k52 k53 k54 k55 k55 k56 k56 k58 k58 k61 k61 k63 k64 k64 k65 k65 k66 (1/2)k67 (1/2)k67 k68 k70 k75 k78
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Referenceb 4 1 2 2 2, 10 2 2 2 8 2 2 2 2 2 2 8 6, 9, 10 10 2 2 2 3 5 8 10 10 3 5 10 7 2 8 8 8 8 8 8 4 4 1 assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed
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Table 1. (continued) Reaction (R185) (R186) (R187) (R188) (R189) (R190) (R191) (R192) (R193) (R194) (R195) (R196) (R197) (R198) (R200) (R201) (R202) (R205) (R206)
ClCO + D ! DCl + CO D + Cl2 ! DCl + Cl Cl + DO2 ! OD + ClO ClCO3 + D ! CO2 + Cl + OD O + HD ! OD + H O + HD ! OH + D OD + OH ! HDO + O O(1D) + DCl ! Cl + OD OD + HCl ! HDO + Cl OH + DCl ! HDO + Cl DO2 + HCl ! HDO + ClO HO2 + DCl ! HDO + ClO O(1D) + DCl ! O + DCl O(1D) + DCl ! D + ClO ClO + HD ! DCl + OH ClO + HD ! HCl + OD O + D + M ! OD + M D + OH + CO2 ! HDO + CO2 H + OD + CO2 ! HDO + CO2
Rate Coefficient k86 k90 k91 k95 (1/2)k104 (1/2)k104 k105 k106 k110 k110 k111 k111 k115 k116 (1/2)k127 (1/2)k127 k128 k130 k130
Referenceb assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed assumed
a The photolytic coefficients (J) are given for a diurnally averaged model at 45°N at the top of the atmosphere (s1). J value is defined by the product of the species’ cross section and the solar spectrum. Two-body and three-body rate coefficients are given in units of cm3 s1 and cm6 s1, respectively. b 1, Bahou et al. [2001]; 2, Baulch et al. [1980]; 3, Baulch et al. [1981]; 4, Cheng et al. [1999]; 5, Lee and Howard [1982]; 6, Prather et al. [1978]; 7, Robie et al. [1990]; 8, Sander et al. [2006]; 9, Trainor et al. [1973]; 10, Yung and DeMore [1982].
to the reference standard, which is set at the lower boundary. Figure 2 shows the ratios of the photolytic coefficients (J values) of HDO/H2O (black curve) and DCl/HCl (gray curve). The ratios decrease with decreasing altitude, demonstrating that their d values can be enhanced accordingly by photolytic processes. The photolytic isotope fractionation studied in this article is caused primarily by self-shielding by the most abundant molecules. The photodissociation of H2O occurs primarily at altitudes between 80 and 110 km and that of HCl at altitudes less than 70 km; the latter is more significant below 95 km in terms of photolysis rate (Figure 3). In the regions below 95 km, the photolysis of HCl provides a source of H atoms that eventually form H2O. As this source favors the formation of H2O over HDO, this results in a depletion of d(HDO). The resulting column-averaged (above the lower boundary of 58 km) d(HDO) is 2%. As a result, HDO, HD, and D are all isotopically depleted; and d(DCl) is enhanced by 25%. The factor of 10 between d(HDO) and d(DCl) is due to higher abundance of H2O relative to HCl. Above 95 km, H2O photolysis is significant and d(HDO) increases with altitude. Including H2O/HDO photolysis enhances its isotopic composition, and consequently, the total d(HDO) above 58 km becomes 1% (Figure 4). [7] Sensitivity of the results to atmospheric transport is shown by the dashed curves in Figures 1 – 4. Though the transport provides a good fit to the observed H2O and HDO profiles, it is negligible in modifying d(HDO). This is caused by a high eddy diffusion coefficient (a few times 105 cm2 s1, or a transport time of 105 sec) that greatly dilutes the isotopic fractionation resulting from H2O/HDO and HCl/DCl photolytic processes (see an explanation of the dilution effect by Liang et al. [2007]). Several other modifications in transport have been tested but none gives a satisfactory explanation to both H2O and HDO profiles. For example, the decreasing mixing ratios of H2O and HDO between 70 and 85 km can be caused by advection and diffusive separation. The former can be prescribed by a downwelling transport (adopted in this paper). The latter
requires a reduction of eddy mixing coefficients (so that molecular diffusion becomes more important than advective transport). The increase above 85 km can be attributed either to upwelling transport (adopted in this paper) or to downwelling of air from a higher region in the atmosphere where H2O/HDO sources are needed. To account for the increase, the required H2O source is 5 108 molecules cm2 s1 (the total photolysis rate of H2O approximately the photolysis rate of HCl above 85 km; Figure 4), a value that is unreasonably high. Furthermore, the increase of the ratio of [HDO]/[H2O] above 70 km observed by Venus Express implies that either H2O photolysis is not small (i.e., at least 0.1 times HCl photolysis) or a hitherto unknown mechanism that transports HDO to the region. Further a general circulation model with the correct physics (e.g., heating between 90 and 120 km) is urgently needed. The laboratory measurements of the photolytic cross sections of H2O/HDO/ HCl/DCl at temperatures (150–200 K) similar to the Venus’ are also required, because of the temperature-dependent nature of PHIFE [e.g., Kaiser et al., 2002; Liang et al., 2004]. [8] Additionally, atomic hydrogen can escape from the atmosphere of Venus. The escape fluxes of H and D are estimated to be 3.5 106 and 3.1 104 atoms cm2 s1, respectively [Gurwell and Yung, 1993]. Higher fluxes determined by Donahue [1999] are also tested, but no noticeable effect is observed. This additional process, that enhances atomic D abundance (i.e., increase D/H ratio) in an air parcel, followed by subsequent chemical reactions of H/D with oxygen compounds favors the production of HDO. However, the escape flux is 2 orders of magnitude lower than the photolysis of H2O/HDO and HCl/DCl. Consequently, the hydrogen escape plays a small role in modifying the isotopic composition of water. This effect is identical to the influx of water described in the previous paragraph. [9] Figure 5 shows the modeled HDO/H2O ratios in the 2-D model. The ratio is relatively uniform in the regions below 95 km and increases above. The increase is caused by the preferential photolysis of H2O over HDO, as de-
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Figure 2. Ratios of the photolytic coefficients (J values) of deuterated and normal molecules. Black and gray curves are for the ratios of HDO/H2O and DCl/HCl, respectively. Line designation is same as that in Figure 1. The solid and dashed curves are indistinguishable in the current plotting scale. scribed above. In addition, we see that the ratio of HDO/ H2O also decreases with increasing latitude, demonstrating that in the current model, transport plays a negligible role in modifying the isotopic composition of water. As in the reference 1-D model, we obtain no kink feature in H2O and HDO mixing ratio profiles observed by Venus Express. The transport was obtained by assuming Newtonian cooling [Lee et al., 2007]. Future work on the transport is required.
4. Discussion [10] In our standard 1-D model, the eddy diffusion coefficient in the upper atmosphere of Venus is 106 cm2 s1, and this corresponds to a mixing time 105 s [see Yung and DeMore, 1982]. The prescribed advection transport has a time constant of 106 s (velocity of 0.4 cm s1 and atmospheric scale height of 4 km), an order of magnitude lower than the eddy mixing time. This confirms that the prescribed advection can explain well the vertical mixing ratio profile of H2O but not the ratio profile of HDO/H2O. We will demonstrate in section 4.2 that the tentatively observed enhancement of HDO relative to H2O [Bertaux et al., 2007] could be explained by the PHIFE of water alone, through hitherto unknown processes that transport the isotopic signature to a lower altitude. Moreover, our model also suggests that in order to explain the observed H2O profile, an upward transport of 0.5 cm s1 between 85 and 95 km and a downward transport of 0.5 cm s1 above 95 km are needed. The existence of downward transport is supported by the temperature anomaly observed by Venus Express [Bertaux et al., 2007] (see section 4.1), but the upward transport remains a puzzle. We note that though the temperature inversion is obtained on the night side at midlatitudes, this inversion probably extends to the polar region on the basis of the analogy with the terrestrial middle atmosphere [see, e.g., Holton et al., 1995, Figure 3]. A 3-D
general circulation model that includes realistic heating is urgently needed to resolve these issues. 4.1. Temperature Anomaly and Descent Rate [11] SPICAV data show a large temperature anomaly 20– 50 K at around 100 km, taken on the night side at low to midlatitudes [see Bertaux et al., 2007, Figure 1]. The authors interpret this as evidence for heating by air subsidence. The associated vertical velocity may be estimated as follows. The thermodynamic equation can be written as follows [Andrews et al., 1987, p. 115]: DT =Dt þ kwT=H ¼ J =Cp ;
ð2Þ
where D/Dt is material derivative, T is temperature, w is vertical velocity, H is scale height, J is diabatic heating rate, k is (g – 1)/g, g is Cp/Cv, and Cp and Cv are heat capacity at constant pressure and volume, respectively. For a CO2 atmosphere, k = 1/4. For vertical advection only, we have DT =Dt ¼ @T =@t þ w@T =@z:
ð3Þ
Substituting (3) into (2) and simplifying, we have @T =@t þ wð@T =@z þ kT =H Þ ¼ J =Cp :
ð4Þ
Since we are interested in dynamical heating, we can set J = 0. Equation (4) allows us to estimate the change of temperature due to dynamical heating DT ¼ wð@T =@z þ kT =H ÞDt:
ð5Þ
The maximum change in T occurs for Dt = trad, where trad is the Newtonian cooling time constant [see, e.g., Goody and Yung, 1989, p. 252]. If Dt trad cooling by radiation becomes important, thereby reducing the effect of dynami-
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Figure 3. The photolysis rates of H2O (black) and HCl (gray). Line designation is same as that in Figure 1.
cal heating. For Venus around 100 km (0.03 hPa), we have the following values for the quantities in equation (5): T ¼ 175K;
ð6Þ
@T =@z ¼ 4K=km;
ð7Þ
H ¼ 4km;
Dt ¼ trad 1 day;
H2 O þ hv ! OH þ H;
ð11Þ
HDO þ hv ! OD þ H;
ð12aÞ
HDO þ hv ! OH þ D:
ð12bÞ
ð8Þ
ð9Þ
where in (9), we have used estimates based on Crisp and Titov [1997]. Substituting (6) – (9) into (5) we have DT ¼ 54w;
summarized in Figure 6. The higher signal-to-noise ratio data are obtained at altitudes between 75 and 85 km (blue symbols). An offset between the solid line and the blue symbols can be accounted for by the mixing ratio of H2O and the ratio of [HDO]/[H2O] set at the lower boundary. Both H2O and HDO are destroyed by photolysis
ð10Þ
where the units for DT and w are K and cm s1, respectively. Thus, if we have downwelling velocity 1 cm s1, then DT is 54 K, which is also the correct order of magnitude to account for the H2O profile, as discussed in the previous section. On the other hand, if we use the descent velocity of 0.43 m s1 suggested by Bertaux et al. [2007], we have DT is 2322 K, which is clearly too large compared with the observations! Thus, the H2O profile offers a potentially valuable clue to the descent rate of air in the mesosphere of Venus. 4.2. Isotopic Enrichment [12] SOIR data show evidence for changes in H2O, HDO and the HDO/H2O ratio. There appears to be a decrease in H2O mixing ratio between 70 km and 85 km, even though the mixing ratio of H2O or HDO could be seriously affected by the uncertainty in CO2 retrieval, as pointed out by Vandaele et al. [2008]. This is explained as follows and
Let x and y be the concentrations of H2O and HDO, respectively. In the absence of production, the loss of x and y is described by dx=dt ¼ J11 x;
ð13Þ
dy=dt ¼ J12 x;
ð14Þ
where J11 and J12 denote the photodissociation coefficients for reactions (11) and (12). Solving these equations, we have xðt Þ ¼ xð0Þ expðJ11 t Þ
ð15Þ
yðt Þ ¼ yð0Þ expðJ12 t Þ;
ð16Þ
and
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Rðt Þ ¼ ½ yðt Þ=xðt Þ=½ yð0Þ=xð0Þ ¼ exp½ðJ11 J12 Þt :
ð17Þ
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Figure 4. Model HDO/H2O corresponding to cases in Figure 1. Data are from Bertaux et al. [2007]. The two curves are indistinguishable below 100 km. Line designation is the same as that in Figure 1.
Let r(t) = x(t)/x(0). From (15), we have rðt Þ ¼ expðJ11 t Þ:
ð18Þ
From (17) and (18), we can derive a simple relation Rðt Þ ¼ ½rðt Þf ;
ð19Þ
f ¼ ðJ11 J12 Þ=J11 :
ð20Þ
where
Referring to the data from SOIR, let us assume that the mixing ratio of H2O at the cloud tops is 1 ppm; at 95 km,
the value has decreased to 0.7 ppm. Therefore, r(t) = 0.7. From our model, J12/J11 = 0.54; thus f = 0.46 (The total column integrated photolysis rates of HDO and H2O are 2.6 107 and 4.8 108 molecules cm2 s1, respectively, giving the averaged ratio of J12/J11 = 0.54. Since the vertical transport time is significantly shorter than the photolytic lifetimes of H2O and HDO, the column averaged J values are relevant to use.). The expected R(t) from (19) is 1.2, which is close to the observed enrichment between 70 and 75 km and 90– 95 km. This simple theory offers a satisfactory explanation of the increase of the HDO/H2O ratio as a result of the preferential destruction of H2O relative to HDO. Figure 6 provides a quantitative prediction for the isotope ratio of water (R) to
Figure 5. Reference case for simulation of the ratio of HDO and H2O by the 2-D model. 7 of 9
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Figure 6. The isotope ratio of water (R) to the fraction of remaining water (r). See text for details.
the fraction of remaining water (r). This relation is consistent with the limited observations currently available. It is desirable to test this relation over a wider range of R and r.
5. Concluding Remarks [13] Simple one-dimensional and two-dimensional chemistry and transport models are used to study the spatial distribution of HCl, H2O, and HDO in the upper atmosphere of Venus. The distributions of these molecules in the atmosphere reflect the influence of chemistry and transport. A large subsidence at around 85 km is needed to explain the observed vertical profile of H2O. The magnitude of the descent rate is consistent with the temperature anomaly observed by SPICAV. Preferential photolysis of H2O over HDO provides a driving force for isotopic enrichment. Laboratory measurements of temperature-dependent cross sections for H2O over HDO are needed. The correct simulation of the distribution of HCl, H2O, and HDO ultimately requires a 3-D model. This paper identifies the most important physical and chemical processes that a model must incorporate. [14] The emphasis of this work is in the region of the atmosphere above 70 km, where the H2O abundance is 1 ppm. As the abundance of H2O in the deep atmosphere is 100 ppm, there is another removal mechanism that is not related to photolysis. It is known that H2SO4 aerosols are formed above the cloud tops. They represent a net sink of SO2 and H2O, thereby explaining their rapid decrease with altitude above the cloud tops. We do not know whether this process can cause fractionation. Laboratory studies are needed to determine the fractionation associated with this chemical dehydration, if any occurs. [15] Acknowledgments. Special thanks are due to J.-L. Bertaux for providing H2O, HDO, HCl, and HF profiles from Venus Express. We thank H. Hartman, N. Heavens. K. F. Li, V. Natraj, C. Parkinson, R. L. Shia, and X. Zhang for critical comments. This research was supported in part by
NSC grant 97-2628-M-001-001 to Academia Sinica and NASA grant NNX07AI63G to the California Institute of Technology.
References Andrews, D. G., J. R. Holton, and C. B. Leovy (1987), Middle Atmosphere Dynamics, 489 pp., Academic Press, Orlando, Fla. Bahou, M., et al. (2001), Absorption cross sections of HCl and DCl at 135 – 232 nanometers: Implications for photodissociation on Venus, Astrophys. J., 559(2), L179 – L182, doi:10.1086/323753. Baulch, D. L., et al. (1980), Evaluated kinetic and photochemical data for atmospheric chemistry, J. Phys. Chem. Ref. Data, 9(2), 295 – 471. Baulch, D. L., J. Duxbury, S. J. Grant, and D. C. Montague (1981), Evaluated kinetic data for high-temperature reactions, Vol. 4—Homogeneous gas-phase reactions of halogen-containing and cyanide-containing species, J. Phys. Chem. Ref. Data, 10(Supplement 1), 1 – 721. Bertaux, J. L., et al. (2007), A warm layer in Venus’ cryosphere and highaltitude measurements of HF, HCl, H2O and HDO, Nature, 450, 646 – 649. Cheng, B. M., et al. (1999), Photo-induced fractionation of water isotopomers in the Martian atmosphere, Geophys. Res. Lett., 26(24), 3657 – 3660, doi:10.1029/1999GL008367. Crisp, D., and D. Titov (1997), The thermal balance of the Venus atmosphere, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher, D. M. Hunten, and R. J. Phillips, pp. 353 – 384, Univ. of Ariz. Press, Tucson, Ariz. Donahue, T. M. (1999), New analysis of hydrogen and deuterium escape from Venus, Icarus, 141(2), 226 – 235, doi:10.1006/icar.1999.6186. Fedorova, A., O. Korablev, A.-C. Vandaele, J.-L. Bertaux, D. Belyaev, A. Mahieux, E. Neefs, W. V. Wilquet, R. Drummond, F. Montmessin, and E. Villard (2008), HDO and H2O vertical distributions and isotopic ratio in the Venus mesosphere by Solar Occultation at Infrared spectrometer on board Venus Express, J. Geophys. Res., 113, E00B22, doi:10.1029/ 2008JE003146. Goody, R. M., and Y. L. Yung (1989), Atmospheric Radiation: Theoretical Basis, Oxford Univ. Press, New York. Gurwell, M. A., and Y. L. Yung (1993), Fractionation of hydrogen and deuterium on Venus due to collisional ejection, Planet. Space Sci., 41(2), 91 – 104, doi:10.1016/0032-0633 (93)90037-3. Holton, J. R., P. H. Haynes, M. E. McIntyre, A. R. Douglass, R. B. Rood, and L. Pfister (1995), Stratosphere-troposphere exchange, Rev. Geophys, 33(4), 403 – 439. Kaiser, J., T. Rockmann, and C. A. M. Brenninkmeijer (2002), Temperature dependence of isotope fractionation in N2O photolysis, Phys. Chem. Chem. Phys., 4(18), 4420 – 4430, doi:10.1039/b204837j. Kasting, J. F., and J. B. Pollack (1983), Loss of water from Venus. Part 1. Hydrodynamic escape of hydrogen, Icarus, 53(3), 479 – 508, doi:10.1016/0019-1035 (83)90212-9. Lee, Y. P., and C. J. Howard (1982), Temperature-dependence of the rateconstant and the branching ratio for the reaction Cl+HO2, J. Chem. Phys., 77(2), 756 – 763, doi:10.1063/1.443892.
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Lee, C., S. R. Lewis, and P. L. Read (2007), Superrotation in a Venus general circulation model, J. Geophys. Res., 112, E04S11, doi:10.1029/ 2006JE002874. Liang, M. C., G. A. Blake, and Y. L. Yung (2004), A semianalytic model for photo-induced isotopic fractionation in simple molecules, J. Geophys. Res., 109, D10308, doi:10.1029/2004JD004539. Liang, M. C., et al. (2005), Meridional transport in the stratosphere of Jupiter, Astrophys. J., 635(2), L177 – L180, doi:10.1086/499624. Liang, M. C., G. A. Blake, B. R. Lewis, and Y. L. Yung (2007), Oxygen isotopic composition of carbon dioxide in the middle atmosphere, Proc. Natl. Acad. Sci. U.S.A., 104(1), 21 – 25, doi:10.1073/pnas.0610009104. Miller, C. E., and Y. L. Yung (2000), Photo-induced isotopic fractionation, J. Geophys. Res., 105(D23), 29,039 – 29,051. Mills, F. P. (1998a), Observations and photochemical modeling of the Venus middle atmosphere, Ph.D. thesis, Calif. Inst. of Technol, Pasadena, Calif. Mills, F. P. (1998b), Thermal infrared spectroscopy of Europa and Callisto, Ph.D. thesis, Calif. Inst. of Technol, Pasadena, Calif. Prather, M. J., J. A. Logan, and M. B. McElroy (1978), Carbon-monoxide in Jupiter’s upper-atmosphere: An extraplanetary source, Astrophys. J., 223(3), 1072 – 1081, doi:10.1086/156340. Robie, D. C., S. Arepalli, N. Presser, T. Kitsopoulos, and R. J. Gordon (1990), The intramolecular kinetic isotope effect for the reaction O(3P)+HD, J. Chem. Phys., 92(12), 7382 – 7393, doi:10.1063/1.458224.
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Sander, S. P., et al. (2006), Chemical kinetics and photochemical data for use in atmospheric studies, Jet Propul. Lab. Pub. 06 – 2, Jet Propul. Lab., Pasadena, Calif. (Available at http://jpldataeval.jpl.nasa.gov/.) Trainor, D. W., D. O. Ham, and F. Kaufman (1973), Gas-phase recombination of hydrogen and deuterium atoms, J. Chem. Phys., 58(10), 4599 – 4609, doi:10.1063/1.1679024. Vandaele, A. C., et al. (2008), Composition of the Venus mesosphere by SOIR on board Venus Express, J. Geophys. Res., 113, E00B23, doi:10.1029/2008JE003140. Yung, Y. L., and W. B. DeMore (1982), Photochemistry of the stratosphere of Venus: Implications for atmospheric evolution, Icarus, 51(2), 199 – 247, doi:10.1016/0019-1035(82)90080-X. Yung, Y. L., M. C. Liang, X. Jiang, C. Lee, B. Bezard, and E. Marcq (2008), Modeling the distribution of OCS in the lower atmosphere of Venus, J. Geophys. Res., doi:10.1029/2008JE003094, in press.
M.-C. Liang, Research Center for Environmental Changes, Academia Sinica, 128 Academia Road, Section 2, Taipei 115, Taiwan. (mcl@rcec. sinica.edu.tw) Y. L. Yung, Division of Geological and Planetary Sciences, California Institute of Technology, MC 170-25, 1200 East California Boulevard, Pasadena, CA 91125, USA.
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Water vapor abundance near the surface of Venus from Venus Express/VIRTIS observations Bruno Be´zard,1 Constantine C. C. Tsang,2 Robert W. Carlson,3 Giuseppe Piccioni,4 Emmanuel Marcq,1 and Pierre Drossart1 Received 13 August 2008; revised 16 December 2008; accepted 13 March 2009; published 1 May 2009.
[1] Nightside observations of the 1.18-mm atmospheric window by the Visible and
Infrared Thermal Imaging Spectrometer (VIRTIS) aboard the Venus Express spacecraft were analyzed to measure and map the water vapor abundance in the lower atmosphere. Thermal emission in this window originates partly from the surface and partly from the first scale height (0–15 km) of the atmosphere. Constraints on the CO2 continuum absorption, which is the dominant source of gaseous opacity in the window, were obtained from the variation of the 1.185-mm intensity with surface elevation. An absorption coefficient of 1 ± 0.4 109 cm1 amagat2 best fits the observed variation. We retrieved a water vapor mole fraction of 44 ± 9 ppm from various selections of VIRTIS spectra in the southern hemisphere, in agreement with previous analyses of the nightside emission. This value is somewhat larger than that previously determined at higher altitudes from the 2.3- and 1.74-mm nightside windows, but the error bars still allow a constant with height H2O mole fraction from the surface up to 40 km. Using the intensity ratio in the two wings of the 1.18-mm window as a proxy, we searched for horizontal variations of the H2O abundance in various VIRTIS observational sequences. We derived stringent upper limits for any possible latitudinal variations on the night side: ±1.5% in the range 60°S–25°N and ±3% for the broader range 80°S–25°N. The lack of detectable latitudinal variations is consistent with a constant with height water profile in the lower atmosphere and probably precludes any strong concentration gradient near the surface. Citation: Be´zard, B., C. C. C. Tsang, R. W. Carlson, G. Piccioni, E. Marcq, and P. Drossart (2009), Water vapor abundance near the surface of Venus from Venus Express/VIRTIS observations, J. Geophys. Res., 114, E00B39, doi:10.1029/2008JE003251.
1. Introduction [2] The near-infrared windows centered at 1.01, 1.10 and 1.18 microns provide a means of probing the lower atmosphere and surface of Venus. In these windows, thermal emission from the surface and the lowest scale height (0 – 15 km) of the atmosphere can leak through the thick sulfuric acid clouds and be detected on the night side of the planet [Taylor et al., 1997]. The 1.10- and 1.18-mm windows are limited on one side by the n 1 + n 2 + n 3 H2O band centered at 1.13 mm and on the other side by CO2 bands (respectively the 2n 1 + 3n 3 band at 1.05 mm and the n 1 + 3n 3 band at 1.21 mm). Imaging and spectroscopic observations of these windows can thus provide information on the water vapor abundance near the surface and its variations. [3] Be´zard and de Bergh [2007] recently reviewed the determinations of the water vapor abundance in Venus’ deep 1 LESIA, Observatoire de Paris, UPMC, Universite´ Paris-Diderot, CNRS, Meudon, France. 2 Atmospheric, Oceanic and Planetary Physics, Department of Physics, University of Oxford, Oxford, UK. 3 Jet Propulsion Laboratory, Pasadena, California, USA. 4 INAF, IASF, Rome, Italy.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003251$09.00
atmosphere. In the 1990s, various ground-based observations of Venus’ night side established that a H2O mole fraction around 30 ± 10 ppm can reproduce the 1.18-mm window’s spectrum [Crisp et al., 1991; Pollack et al., 1993; de Bergh et al., 1995]. From spectroimaging data covering a broad range in surface altitude, Meadows and Crisp [1996] derived a slightly larger H2O mole fraction, 45 ± 10 ppm. They also argued that the temperature lapse rate in the lowest 6 km is shallower than that measured in situ [e.g., Seiff, 1983] and used in earlier analyses of nightside observations. Reanalyzing Venera 11, 13 and 14 optical spectra, Ignatiev et al. [1997] concluded that the H2O mixing ratio lies between 20 and 40 ppm in the altitude range 5 – 60 km. Below 5 km, the quality of the Venera data is worse but they seem to indicate an increase of the H2O mixing ratio up to 50– 70 ppm at the surface. All these results show that significant uncertainties remain on the water abundance close to the surface. In particular, analysis of near-infrared windows is hampered by our poor knowledge of the CO2 ‘‘continuum’’ opacity at high pressures, high temperatures, and long path lengths, arising from collisioninduced bands and extreme far wings of allowed CO2 bands. [4] Drossart et al. [1993] searched for horizontal variations of the H2O abundance using Galileo/NIMS data
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Figure 1. Two averages of 135 Venus’ nightside spectra recorded during session 390– 13 over Alpha Regio with a mean altitude of 1.30 km (thick line) and about 5° to the east at a mean altitude of 0.47 km (thin line). Emission angle is 45° for both selections. Spectra have been shifted by 7 nm to correct for a slight wavelength calibration error. The location of the main CO2 and H2O bands that limit the nearinfrared windows is shown along with the position of the O2 airglow emission. recorded over a strip on the night side extending from 60°S to 40°N. Analyzing the 1.18-mm window, they concluded that the water abundance in the 0 –15 km range did not vary by more than 20% over the limited area covered. No other searches for horizontal or temporal variations of water vapor near the surface have been published. [5] The Venus Express spacecraft, orbiting Venus since April 2006, is currently monitoring the dynamics and composition of Venus’ atmosphere. In particular, the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) is repeatedly mapping Venus’ nightside emission through its infrared imaging channel (M-IR), giving access to the composition of the deep atmosphere over large horizontal and temporal scales. We present here an investigation of the water vapor abundance near the surface of Venus using various VIRTIS-M observational sets. We focused on the 1.18-mm window, which is the most favorable, being 3 times more intense than the 1.10-mm one. In a first step, we derived constraints on the CO2 so-called continuum opacity by investigating the variation of the 1.18-mm emission as a function of the surface elevation. We then determined the H2O mixing ratio that best reproduces observed spectra. Finally, we analyzed several VIRTIS-M spectral images to search for possible weak horizontal abundance variations, using intensity ratios in the 1.18-mm window as a proxy.
2. VIRTIS Observations [6] VIRTIS is a dual spectrometer comprising two subinstruments with their own telescope: an imaging spectrometer (VIRTIS-M) with visible and infrared channels and a high-resolution spectrometer (VIRTIS-H) limited to the interval 2 – 5 mm [Piccioni et al., 2007]. The infrared channel of VIRTIS-M covers the range 1.05 – 5.2 mm with a spectral sampling of 9.5 nm. The spatial sampling is
0.25 mrad and the instantaneous field of view (FOV) is 0.25 64 mrad, with 256 pixels along the slit. Scanning in the direction across the slit with 256 step positions at 0.25 mrad per step yields a 256 256 image with a 64 64 mrad FOV. [7] The VIRTIS-M spectral cubes used in this study are listed in Table 1. We selected observational sequences offering all together a good latitudinal coverage from south pole to low northern latitudes (10 – 20°N). They usually correspond to off-pericenter observations (Science case 2) except for sequence 093– 01, which corresponds to Science case 3 (global spectroimaging from apocenter). The Venus Express science planning is described by Titov et al. [2006]. The integration time varies from 3 to 8 s, ensuring a good signal-to-noise ratio on the night side. The sequences marked in bold incorporate regions of relatively high elevation (Alpha Regio, Imdr Regio, Themis Regio) along with regions of low elevation, and have been used to constrain the CO2 continuum opacity (see section 4.2). Figure 1 shows an example of VIRTIS spectra averaged over two small areas observed in session 390 – 13. [8] A small amount of scattered sunlight is visible in the spectra at short wavelengths (Figure 1). To remove it from the nightside spectra, we used the intensity at 1.061 mm (Spectel 4) and 1.232 mm (Spectel 22), two wavelengths at which no nightside emission is expected from synthetic models. We assumed a linear spectral variation for this scattered component between these two wavelengths. The so-calculated scattered light also reproduces the residual intensity at 1.356 mm (Spectel 35), another wavelength at which models predict no nightside emission. [9] We found that the wavelength scale of the observations was slightly shifted from that given in the spectral cubes after the latest calibration of the VIRTIS VEX archive (v2.1, 18 March 2008). In the sequences selected, this shift
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Table 1. VIRTIS Nightside Sequencesa Sessionb 093 – 01 157 – 03 390 – 09 390 – 13 392 – 05 392 – 06 571 – 08 571 – 11 577 – 06 577 – 08 577 – 10 579 – 02 579 – 06 579 – 08 579 – 10 579 – 12
Date 22 25 16 16 17 17 12 12 18 18 18 20 20 20 20 20
Jul 2006 Sep 2006 May 2007 May 2007 May 2007 May 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007 Nov 2007
Science Case
Exposure Time (sec)
Maximum and Minimum Latitudes
3 2 2 2 2 2 2 2 2 2 2 2 2 2 2 2
3.3 3.3 8 8 8 8 3 3 8 8 8 8 8 8 8 8
7; 61 8; 44 9; 68 17; 56 34; 85 29; 82 5; 61 8; 70 10; 56 14; 55 15; 50 46; 90 9; 57 14; 55 16; 50 23; 31
Easternmost and Westernmost Longitudes 178; 331; 323; 331; 328; 337; 199; 245; 279; 223; 186; 178; 276; 228; 192; 229;
247 6 29 23 119 114 269 311 327 281 229 340 333 287 243 265
Comment includes Imdr Regio includes Alpha Regio includes Alpha Regio includes Imdr Regio includes Themis Regio
a
Visible and Infrared Thermal Imaging Spectrometer. Sessions in bold are those used to derive the CO2 continuum opacity (section 4.2).
b
varies between approximately 0.75 and 0.95 time the spectel size (9.5 nm). It varies reproducibly along the 256-pixel slit by 0.1 pixel, reaching a maximum around Pixel 100 (see section 5 and Figure 7). We considered this spectral shift as a free parameter in the fitting procedure used to derive the water vapor abundance.
assumed the following sublorentzian lineshape (with a cutoff at 180 cm1 from line center):
3. Radiative Transfer Model
where Ds is the distance from line center expressed in cm1 and c is the factor by which we multiply the Voigt lineshape. This c factor was found to best reproduce the low-frequency side of the 1.18-mm window, dominated by CO2 absorption, as observed in high-resolution groundbased spectra [de Bergh et al., 1995]. Various analyses of the Venus nightside windows have shown that an additional continuum opacity, likely due to collision-induced CO2 bands and/or extreme far wings of strong allowed CO2 bands, is a major contributor to the gas opacity. Its strength may be relatively well determined in the 2.3- and 1.74-mm windows by fitting the amplitude of allowed CO2 lines in high-resolution spectra [e.g., Be´zard et al., 1990; Pollack et al., 1993; de Bergh et al., 1995]. This is not the case for the 1.18- and 1.1-mm windows where no marked CO2 feature is present. Fortunately, it is possible to take advantage of the transparency of these two windows and thus of their sensitivity to the surface elevation to determine this ‘‘continuum’’ opacity, as shown in section 4. We parameterize it with a constant binary absorption coefficient, expressed in cm1 amagat2. [12] For each spectral selection we analyzed, we used the mean outgoing intensity from our model (i.e., the upward flux divided by p) and rescaled it by a factor f to reproduce the observed 1.18-mm intensity. This factor accounts for the variation in emission angle and cloud opacity from one selection to another. This simplified approach is valid because these two parameters are spectrally neutral and uniformly modulate the outgoing intensity. To check this, we calculated the radiance ratios of spectra at various emission angles up to 80° to that at zero emission angle. Figure 2 shows that these ratios are quasi-constant across the 1.18-mm window, with a variation less than 1%. The water vapor signature is thus fully preserved in changes of
[10] Synthetic spectra were calculated using a line-by-line radiative transfer model with scattering originally developed to analyze high-resolution spectra of Venus’ night side [Be´zard et al., 1990; de Bergh et al., 1995]. We improved over these earlier works by employing the Discrete Ordinates Radiative Transfer Program for a MultiLayered PlaneParallel Medium (DISORT) algorithm [Stamnes et al., 1988] with eight streams to solve the equation of transfer in place of a Delta-Eddington adding algorithm. A HenyeyGreenstein phase function was used for particle scattering. The temperature profile, assumed to be horizontally uniform, is based on entry probe data at low latitudes, as compiled by Seiff [1983, Table A1]. We did not update these data, the VIRA-2 profile [Moroz and Zasova, 1997] being the same as the VIRA one in the lower atmosphere. The CO2 mole fraction is 0.965 and the H2O profile is assumed constant with height below the clouds. Rayleigh scattering by CO2 is included. Surface emissivity was arbitrarily set to 0.95. We used the simplified cloud model described by Crisp [1986] that extends from 30 to 80 km. Extinction efficiencies, single scattering albedos and asymmetry parameters were generated layer by layer by calculating Mie scattering by four modes of 75% H2SO4 particles. [11] For CO2 line opacity, we used the high-T database presented by Pollack et al. [1993], while the H2O line parameters come from Geisa 97 [Jacquinet-Husson et al., 1999]. The CO2 self-broadened half width was taken as 0.1 cm1 atm1 and varying as T0.75. For CO2 – broadened H2O line half widths, we used a routine provided by R.H. Tipping and R. Freedman (private communication, 2000). We used a Voigt profile for the H2O lineshape with a cutoff at 120 cm1 from line center. For CO2 lines, we
Ds < 3 : c ¼ 1 3 < Ds < 60 : c ¼ 1:051 expðDs=60Þ 60 < Ds : c ¼ 0:6671 expðDs=110Þ;
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Figure 2. Ratios of synthetic spectra calculated at 17-nm resolution for different emission angles (solid lines), twice and half the optical depth of the lower cloud layer (dashed lines), and for surface elevation of 1 and 2 km (lines with triangles). In all cases, the reference spectrum corresponds to the nominal cloud model described in the text, a zero surface elevation, and a zero emission angle. the emission angle and the mean intensity throughout the window is proportional to the intensity at any emission angle. This occurs because the clouds are optically thick and located well above the line-forming region. In Figure 2, we also show the radiance ratios of models in which the cloud optical depth in the middle and lower cloud layers (below 57 km) is equal to half and twice that in the nominal model. A negligible variation of these ratios across the window is found (less than 1%). Cloud scattering therefore acts a gray attenuator (or a scaling factor) and does not modify the spectral shape of the window because it takes place well above the line-forming region at 1.18 mm with negligible thermal emission by the particles themselves. This also means that our water vapor retrievals are not sensitive to uncertainties in our cloud model.
4. Determination of the Water Vapor Abundance 4.1. Spectral Resolution of VIRTIS-M Observations [13] A precise knowledge of the spectral resolution of the VIRTIS spectra is important for an accurate determination of the H2O abundance, all the more as the wavelength calibration of the data is not known to better than about half a spectel. Water vapor absorption affects the width of the 1.18-mm window so that an error in the spectral resolution can be offset to first order by a change in the H2O abundance and a slight variation in the pixel registration. The spectral profile of the VIRTIS-M infrared channel, around the center of the slit, has been measured before launch. The profile could be fit with a Gaussian having a full width at half maximum (FWHM) of 11.0 ± 1.2 nm. However, synthetic spectra convolved with this profile do not match the VIRTIS Venus nightside spectra, as evidenced at 1.74 mm, the narrowest window. To determine the actual resolution of the observations, we used this window and compared VIRTIS spectra to high-resolution ground-based
spectra recorded at the Canada-France-Hawaii telescope (CFHT) by one of us (BB), Catherine de Bergh, Dave Crisp and Jean-Pierre Maillard [Taylor et al., 1997, Figure 4]. We convolved the latter spectrum with Gaussian functions having different FWHMs and determined the best value through a least squares fit. Figure 3 shows the result for the Alpha Regio spectrum of Figure 1: The CFHT spectrum convolved at 11 nm resolution is too narrow compared to the VIRTIS spectrum and does not produce enough emission in the wings of the window whereas a FWHM of 17 nm yields a very good fit. We found that, among the VIRTIS sessions marked in bold in Table 1, the FWHM of the 1.74-mm window varies between 15 and 19 nm. We used the so-inferred resolution to analyze the 1.18-mm window. The actual resolution of the VIRTIS-M infrared spectra thus appears to be at least 50% larger than measured in the laboratory owing to different temperatures of the spectrometer (135 K for the laboratory measurement versus around 150– 160 K in orbit). 4.2. Determination of the CO2 Continuum Absorption [14] As discussed in section 3, the presence of an unknown continuum opacity in the windows, likely due to CO2 collision-induced bands or extreme far wings of strong allowed bands, adds some uncertainty in the composition retrievals. To constrain this quantity, we analyzed VIRTIS 1.18-mm spectra taken over regions of different elevations during the same session. The contrast between high and low elevation spectra, due to differences in surface temperatures, is sensitive to the gas opacity, e.g., a larger continuum absorption increases the atmospheric opacity, attenuates more strongly the surface emission from lower elevations, and thus diminishes the contrast. In five different sessions, we selected spectra over areas at relatively high elevation (Alpha Regio, Themis Regio or Imdr Regio) and at low elevation. We used the altimetry derived from Magellan and
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Figure 3. The 1.74-mm VIRTIS spectrum of Alpha Regio in Figure 1 (squares) is compared with CanadaFrance-Hawaii telescope (CFHT) spectra convolved with a Gaussian function having full width at half maximum (FWHMs) of 11 nm (dashed line) and 17 nm (solid line). In each case, a multiplying factor f that minimizes the sum of squares of residuals is applied to the CFHT data, a 0.0007 W sr1 m2 mm1 is added to account for residual scattered sunlight in the instrument, and a wavelength shift lshift = 6.8 nm is applied to the VIRTIS spectrum. smoothed it with a boxcar average of width 80– 100 km to account for blurring by the overlying cloud layers [e.g, Moroz, 2002]. The spectral samples, listed in Table 2, comprise 12 to 285 spectra. [15] Besides elevation, the intensity of the emission at 1.18 micron is also dependent on the emission angle and the cloud opacity. We then corrected the spectra to zero emission angle using the limb-darkening model derived by Carlson et al. (manuscript in preparation, 2008) and we used the peak intensity in the windows at 1.27, 1.31 and 1.74 microns, not probing down to the surface, as an indicator of the cloudiness. The 1.31-mm window (Spectel 30) is in principle best suited for this purpose [Meadows and Crisp, 1996] but it is weak and has a relatively low
signal-to-noise ratio. The 1.28-mm intensity at Spectel 27 is stronger but slightly contaminated by the 1.27-mm O2 airglow peaking at Spectel 26. The 1.74-mm window (Spectel 76) has a very good signal-to-noise ratio but, being at much longer wavelengths, it is not a perfect indicator of the cloud opacity at 1.18 mm: spectra with similar 1.74-mm radiances can differ at 1.18 mm owing to different particle size characteristics as was investigated from Galileo/NIMS data in the case of the 2.3- and 1.74-mm windows by Carlson et al. [1993]. To select low elevation regions having similar cloudiness as the high elevation regions, we used a combination of these three windows and empirically imposed intensity variations less than 4% at 1.28 mm, 6% at 1.31 mm and 15% at 1.74 mm. In each session, we
Table 2. VIRTIS Nightside Spectra Used to Determine the CO2 Continuum Opacity and H2O Mole Fraction a
Session
Pixel Range
093 – 01 093 – 01 093 – 01 093 – 01 157 – 03 157 – 03 157 – 03 390 – 13 390 – 13 390 – 13 571 – 08 571 – 08 571 – 11 571 – 11 571 – 11
120 – 122; 70 – 73 85 – 92; 73 – 79 134 – 143; 95 – 103 144 – 158; 61 – 66 64 – 71; 214 – 220 88 – 96; 177 – 183 19 – 27; 88 – 96 112 – 130; 149 – 163 141 – 159; 126 – 140 69 – 80; 147 – 163 218 – 220; 78 – 81 248 – 254; 58 – 65 93 – 97; 137 – 145 88 – 98; 33 – 43 131 – 139; 138 – 159
Mean Latitude and Longitude 46.4; 43.4; 41.4; 48.8; 26.0; 26.5; 36.9; 26.6; 25.9; 31.7; 46.2; 46.8; 40.8; 51.5; 43.5;
214.8 220.7 208.7 208.2 359.2 355.0 351.9 359.9 354.8 4.2 214.6 206.8 277.4 291.1 270.2
Mean Local Time
Mean Altitude (km)
Mean Emission Angle (deg)
23.10 22.68 23.48 23.51 2.76 3.04 3.24 21.38 21.04 21.66 22.61 22.09 21.22 20.30 21.69
2.58 0.03 1.35 0.14 1.68 0.37 1.11 1.30 0.47 0.13 2.68 0.02 2.06 0.07 0.17
36 38 43 35 28 26 15 45 45 41 32 33 37 32 32
a
Along the slit; across the slit. Assuming a CO2 continuum absorption of 1.1 109 cm1 amagat2.
b
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H2O Mole Fractionb (ppm) 42 43 41 39 39 41 47 43 42 45 49 51 49 54 48
± ± ± ± ± ± ± ± ± ± ± ± ± ± ±
3 3 3 3 4 4 5 3 3 3 4 4 4 5 5
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Figure 4. Relative variations of the intensity at 1.185 mm as a function of surface elevation for the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) spectral selections in Table 2 (symbols). Model variations for different continuum opacity (0.5, 1.0, and 1.5 109 cm1 amagat2) are shown for comparison. chose a region around 0 km altitude and, if possible, one at a lower elevation. Figure 1 shows two averaged spectra from session 390– 13, one at 1.30 km and the other one at 0.47 km. One clearly sees that the intensities at 1.1 and 1.18 mm are larger for the low elevation spectrum whereas the intensities in the 1.27, 1.31 and 1.74-mm windows are similar in the two spectra, which indicates similar cloudiness. [16] Figure 4 shows the relative intensity observed at Spectel 17 (1.185 mm) for the highest and lowest elevations of each session with respect to that at 0 km altitude in the same session. We estimate that each datum point bears an uncertainty of ±0.02 owing to the cloud opacity. A least square fit to these data indicates a variation of the intensity of 5.7 ± 0.4% per km. Comparison with model calculations indicates that a constant absorption coefficient of 1 ± 0.4 109 cm1 amagat2 best reproduces the observed variation. The corresponding optical depth at 0 km is 0.92, larger than that due to H2O lines (0.36 for a mole fraction of 40 ppm) and CO2 lines (0.68) at 1.182 mm and 17-nm resolution, which shows that this continuum opacity is an important contributor to the gaseous absorption in this window. The Rayleigh scattering optical depth is 0.65. We checked that changing the surface emissivity (from 0.95 to 0.80) does not significantly affect the derived continuum absorption coefficient. 4.3. Determination of the H2O Mole Fraction [17] To determine the water vapor abundance, we fitted the spectra in Table 2 with synthetic models calculated with the above determined continuum opacity and various H2O mole fractions. For each value of H2O, we determined the values of the spectral shift lshift and scaling factor f (related to cloud opacity and emission angle) that provide the best fit of the 1.18-mm VIRTIS spectrum and calculate the residuals of the fit. We then retained the value of the H2O mole fraction that minimizes the sum of squares of these resid-
uals. In Figure 5, the Alpha Regio spectrum (Figure 1) is compared to synthetic spectra calculated with 30, 40 and 55 ppm of H2O, which shows the strong effect of water vapor absorption on the spectra. Figure 6 shows that the two spectra of Figure 1 that have different surface elevations are well fitted with a continuum absorption coefficient of 1 109 cm1 amagat2 and a water mole fraction of 40 ppm (the same spectral shift and scaling factor are applied to both spectra). [18] The H2O mole fraction that best reproduces the spectral selections, assuming a continuum absorption coefficient of 1.1 109 cm1 amagat2, is indicated in the last column of Table 2 with the 1 SD error bar derived from the residuals of the fit. The weighted mean value is 45 ppm (44 ppm if a continuum absorption coefficient of 1 109 cm1 amagat2 is used) and the standard deviation of the sample is 5 ppm. The ±40% uncertainty on the CO2 continuum absorption adds a ±3 ppm uncertainty on the H2O mole fraction and the estimated ±10% uncertainty on the VIRTIS-M spectral resolution at 1.18 mm adds another ±7 ppm. Combining quadratically all these uncertainties, we conclude that the H2O mole fraction in the lowest scale height of the atmosphere is 44 ± 9 ppm. Note that this uncertainty does not take into account any possible inaccuracies or incompleteness in the H2O spectral database used here (Geisa 97).
5. Search for Horizontal Variations [19] We searched for possible variations of the H2O abundance in the sequences listed in Table 1. Besides the H2O abundance, the intensity of the 1.18-mm window emission is sensitive to the cloud optical depth, emission angle and surface elevation (Figure 2), all of which vary spatially. It is thus necessary to correct the images for the variations in these parameters or to work on quantities that
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Figure 5. The Alpha Regio spectrum of Figure 1 is compared to synthetic spectra with H2O mole fractions of 30, 40, and 55 ppm. The spectral shift lshift = 7.9 nm and scale factor f = 0.78 are those that yield the best fit of the data with the 40-ppm calculation. are least sensitive to them. The effect of the first two parameters is spectrally neutral across the window (Figure 2) and can therefore be eliminated by ratioing the intensity at two wavelengths. A complication arises from the fact that the spectral registration is not constant along the slit as mentioned above and that the spectral resolution may also slightly vary. The first step is thus to determine the spectral shift for each pixel of the images. This was done by determining the peak of the 1.18-mm window from a parabolic fit of the intensity at spectels 16, 17 and 18 and assigning it to 1.182 mm, as determined from synthetic
calculations (Figure 6). The result is shown for a particular sequence (579 – 08) in Figure 7. In all sequences, the spectral shift increases by 1 nm (0.1 spectel) along the slit from Pixel 6 up to Pixels 80 –100 and slightly decreases beyond. The variation of the spectral shift in the scan direction, i.e., during the acquisition, is negligible for this sequence (Figure 7) and all those investigated here. [20] We first considered the ratio R1 of the intensity at 1.169 mm to that at the peak of the window (1.182 mm) to map water vapor variations. The first wavelength lies in the wing of the window that is dominated by H2O absorption.
Figure 6. The two spectra of Figure 1 (session 390 –13) at elevations of 1.30 and 0.47 km are compared with our best fit model that includes a constant absorption coefficient of 1 109 cm1 amagat2 and a H2O mole fraction of 40 ppm. The same spectral shift and scale factor as in Figure 5 are used for both spectra. 7 of 12
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Figure 7. (top) Map of the 1.182-mm intensity (W sr1 m2 mm1) in sequence 579 – 08 and of the spectral shift, expressed as a fraction of a spectel, with respect to the VIRTIS/Venus Express archive calibration (v2.1). (bottom) Plots of the shift as a function of detector line number along the slit for two different frames (i.e., image lines). The intensity at 1.182 mm is derived from the above mentioned parabolic fit and that at 1.169 mm is interpolated linearly between the two adjacent spectels. Synthetic calcu-
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lations show that a 10% variation of the H2O mixing ratio yields a 3.5% variation of this ratio. Because this ratio is potentially very sensitive to remaining uncertainties in the spectral registration, we considered a second parameter (W), which is the FWHM of the window. Calculations indicate that a 10% variation of the H2O mixing ratio yields a 2.3% variation of this quantity. The maps of these two quantities are shown in Figure 8 for sequence 579 – 08. Clearly, the effect of cloud opacity and emission angle variations is very well removed as attested by comparing to the 1.182-mm image. However, features remain that are not due to water vapor variations. First, there is a slight increase of the two parameters R1 and W with increasing pixel number along the slit (from left to right in Figure 8) for all image lines. This probably reflects a slight increase of the spectral width of the instrumental function from the beginning (small detector line numbers) to the end (large detector line numbers) of the 256-pixel slit. Second, topographic features are clearly seen in the R1 and W images and actually a very good correlation is observed with the altimetry map also shown in Figure 8. This occurs because altitude variations affect more the peak of the window, where the gas opacity is at minimum, than its wings. A higher altitude region is therefore characterized by a larger value of R1 and a larger FWHM as the 1.182-mm intensity is more strongly reduced than those at half maximum in the wings. Consequently, we defined a third parameter R2 to minimize the influence of the topography and of the spectral resolution, that is the ratio of the intensity in the CO2 wing at 1.1945 mm to that in the H2O-dominated wing at 1.169 mm. These two wavelengths exhibit similar intensities and probe similar atmospheric levels so that their intensity ratio is not too sensitive to the surface altitude. Synthetic calculations indicate that this ratio shows little sensitivity to small variations of the spectral resolution and that a 10% variation in the H2O mixing ratio causes a 5% variation of this intensity ratio R2. [21] The map of the R2 ratio, shown in Figure 8, is bland except for the hot pixels, which confirms that all the effects above mentioned have been removed with a very good accuracy. Figure 9 shows cuts through this image with peakto-peak variations of about 3%, attributable to noise and a
Figure 8. Maps of the 1.182-mm intensity I (W sr1 m2 mm1), smoothed surface elevation Z (km), ratio of the intensity at 1.169 and 1.182 mm (R1), FWHM of the window (spectels), and ratio of the intensity at 1.1945 and 1.169 mm (R2) for sequence 579 –08. 8 of 12
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from the day side and we set an envelope of ±1% for any possible latitudinal variation in the whole range 80°S – 25°N. These upper limits translate into ±1% (±0.5 ppm) for the H2O mixing ratio in the range 60°S – 25°N and ±2% (±1 ppm) for the broader range 80°S–25°N.
6. Discussion
Figure 9. Image of the 1.1945- to 1.169-mm intensity ratio (R2) shown in Figure 8, and cuts of this image at line (frame) 70 and at column (detector line) 20. residual altitude signature at the 1% level. We have calculated the median of R2 in latitude boxes 5° wide (Figure 10) and found no variations beyond ±0.3% from the mean from 60°S to 15°N for this sequence 579– 08. The same analysis was performed for all sequences in Table 1 (examples shown in Figure 10). Data southward of 80°S cannot be used owing to a large amount of scattered light from the dayside. Note that for some sequences, having a lower spectral shift than in most cases, we used other wavelengths to calculate the R2 ratio (namely 1.1905 and 1.1685 mm), which allows us to better correct the above mentioned effects. We do not see any significant variations of the R2 intensity ratio with latitude in any sequence and we place an upper limit of ±0.5% for the range 60°S – 25°N. The data poleward of 60°S are somewhat affected by scattered light
[22] The absorption coefficient we inferred for the additional continuum opacity in the 1.18-mm window (1 ± 0.4 109 cm1 amagat2) agrees with that derived independently by Carlson et al. (manuscript in preparation, 2008) from VIRTIS and Galileo/NIMS data. Both analyses assume a temperature lapse rate from the VIRA model [Seiff, 1983] in all investigated regions. If it is shallower (and thus more stable) than in the VIRA model, as argued by Meadows and Crisp [1996], a smaller continuum opacity would be needed. Also, our analysis relies on the assumption of constant surface emissivity at 1.18 mm over the regions we investigated. If, for example, the surface emissivity instead decreases with elevation in the near-infrared as it generally does in Magellan radar images, a larger continuum opacity would be needed to reproduce the observed intensity variation. To our knowledge, we provide here the first determination of this continuum opacity, which is very difficult to measure in the laboratory as it requires very large optical path lengths and high temperatures. [23] Having constrained this important source of opacity, we could determine the H2O mole fraction with a relatively good accuracy. The result, 44 ± 9 ppm, pertains to the 0– 15 km range. It agrees with most previous determinations as reviewed by Be´zard and de Bergh [2007], in particular with that of Meadows and Crisp [1996] (45 ± 10 ppm). We do not have any vertically resolved information given the small range of surface elevation in the data sets we analyzed (3 km) and therefore we cannot discriminate between the somewhat different conclusions of Meadows and Crisp [1996] and Ignatiev et al. [1997]. From ground-based observations of the 1.0, 1.1 and 1.18-mm windows and using the topography for altitude resolution, Meadows and Crisp concluded that the H2O mole fraction was constant (within ±10%) in the lowest scale height (0 – 15 km) at a value of 45 ± 10 ppm. In contrast, from a reanalysis of the Venera optical spectra, Ignatiev et al. [1997] argued that the H2O mole fraction probably increases from about 20 ppm at 10– 20 km to 50 –70 ppm below 5 km. Both distributions yield column-averaged mixing ratios (45 and 37 ppm respectively) that agree with our result within error bars. [24] It is important to note that our determination was obtained using the Geisa 97 database for the H2O absorption. Given the large optical paths and high temperatures involved, it is possible that weak high-energy lines, not included in Geisa (or Hitran), provide a significant contribution in the H2O wing of the 1.18-mm window. If this was the case, the mole fraction needed to reproduce the 1.18-mm observations would probably be lower than the 44 ppm we derived. [25] The 2.3-mm nightside window gives the opportunity to measure the water abundance at higher altitudes, in the range 30– 45 km. A recent analysis of VIRTIS-H spectra indicates a mixing ratio of 31 ± 2 ppm [Marcq et al., 2008], where the error bars derive from the residuals of the fits and
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Figure 10. Variations of the intensity ratio R2 with latitude in sequences 579– 08, 577– 06, 577 – 08, and 577– 10. Images of these sequences at 1.182 mm are also shown. do not account for systematic errors such as in the CO2 continuum absorption. This result is consistent with previous ground-based observations, which indicated 30 ± 10 ppm (see review by Taylor et al. [1997]). Emission in the 1.74-mm window is also sensitive to water vapor absorption and originates from 15 to 25 km, a region intermediate between those probed at 2.3 and 1.18 mm. Ground-based observations of this window again indicate a mole fraction of 30 ± 10 ppm [Pollack et al., 1993; de Bergh et al., 1995]. The 2.3- and 1.74-mm results preclude the large abundance (67 ppm) derived from the Pioneer Venus mass spectrometer measurements [Donahue and Hodges, 1992]. [26] Taken at face value, our H2O mole fraction is about 50% larger than that derived at 2.3 and 1.74 mm, which suggests an increase of the water abundance in the lowest scale height of the atmosphere. However, given the error bars, we feel that a constant mole fraction around 30 –40 ppm from below the clouds down to the surface is still consistent with all analyses of Venus’ nightside emission. A uniform profile agrees with chemical models, which do not predict a significant variation of the H2O mole fraction between 0 and 38 km [Krasnopolsky, 2007]. Below the clouds, the only significant sink for H2O is formation of H2SO4 vapor above 38 km. At 40 km, the H2O mole fraction is depleted by about 2 ppm and at 45 km by about 7 ppm [Krasnopolsky, 2007]. The 2.3-mm window has some weak sensitivity to the 40– 45 km region and high spectral resolution observations might be able to measure this decrease with height. If real, the strong increase of the water mole fraction below 5 km suggested by Ignatiev et al. [1997] would imply a source at or near the surface and a sink above 5 km that are not predicted by existing chemical models and remain to be identified. [27] We have obtained very stringent upper limits for any variation of the H2O mole fraction at 0– 15 km with latitude in several VIRTIS-M sequences: ±0.5 ppm in the range 60°S – 25°N and ±1 ppm if we extend the range to 80°S– 25°N. Drossart et al. [1993] previously searched for H2O variations using the same 1.18-mm window in the Galileo/ NIMS data. This study was limited to three narrow north to south strips covering small areas between 40°S and 50°N.
No variation exceeding ±20% (±6ppm) was found in this latitude range (±10% for the smaller range 30°S – 30°N). Marcq et al. [2006] searched for horizontal variations of H2O in their NASA/IRTF observations of the 2.3-mm nightside window, probing the 30– 45 km region. They did not see any beyond ±15% (±4 ppm) in the latitude range 40°S – 40°N. In contrast, CO shows a 30 – 40% increase from the equator to 60°S at 36 km and an anticorrelated decrease of OCS at 33 km [Marcq et al., 2006, 2008; Tsang et al., 2008]. These variations are interpreted as the signature of the large-scale circulation, with upwelling at the equator and downwelling at high latitudes, in the presence of vertical gradients of CO and OCS due to chemical sources and sinks. The fact that we do not see any variations of the H2O abundance in the first scale height of the atmosphere is fully consistent with the constancy of the vertical profile with height below 40 km and the absence of significant sources and sinks. A contrario, our result probably precludes the factor of 2 –3 increase of the H2O mole fraction from 10 to 0 km advocated by Ignatiev et al. [1997]: General Circulation Model (GCM) calculations by Lebonnois et al. (manuscript in preparation, 2008) suggest an alternance of upward and downward motions from equator to pole in the lowest 20 km, which would probably induce some measurable latitude variation of the H2O abundance we determined, with larger (smaller) abundances in upwelling (downwelling) regions.
7. Conclusions [28] Observations of the night side of Venus by the Venus Express/VIRTIS-M instrument in the 1.18-mm atmospheric window have been used to measure and map the distribution of H2O near the surface of Venus. We have first determined the continuum opacity present in this window from the variation of the peak intensity with surface elevation. This opacity likely results from CO2 collision-induced bands and far wings of strong CO2 bands outside the window (beyond 180 cm1, which is the cutoff we used here), as is the case at 2.3 mm [Tonkov et al., 1996]. The absorption coefficient we inferred is 1 ± 0.4 109 cm1 amagat2. At the
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VIRTIS-M resolution (17 nm), the corresponding optical depth is similar to that due to molecular bands of CO2 and H2O. Up to now, CO2 continuum opacity has been measured only in the 2.3-mm window and at room temperature [Tonkov et al., 1996]. Laboratory measurements in the 1.18and 1.1-mm windows at temperatures relevant to Venus’ deep atmosphere would be extremely useful but are very difficult owing to the huge CO2 path lengths and high temperatures (600 – 730 K) required. [29] Analyzing spectral selections in various areas of Venus’ southern hemisphere, we derived from radiative transfer calculations a water vapor mole fraction of 44 ± 9 ppm, which pertains to the altitude range 0 to 15 km. The quoted error bars do not include possible errors or incompleteness of the H2O spectroscopic data used here (Geisa 97). It is highly desirable to assess whether the Geisa or Hitran databases are sufficient to model the H2O absorption in the 1.18-mm window. If weak hot lines, not included in these databases, contribute significantly to the opacity, the H2O mole fraction needed to reproduce the VIRTIS data would probably be lower than inferred here. Besides this caveat, the main source of uncertainty lies in the spectral resolution of the VIRTIS-M observations that varies from orbit to orbit and was determined for each sequence using the narrow 1.74-mm window. On the other hand, the VIRTIS data have the advantage of an excellent spatial resolution, which permits to obtain spectra at welldefined surface elevations. The limited range of surface elevations (0.5 to 2.7 km) covered by the present analysis did not allow us to retrieve a vertical profile for water vapor following Meadows and Crisp’s [1996] approach. An extension of the Venus Express mission should provide observations of higher terrains, such as Aphrodite Terra, which could constrain the H2O profile in the 0 – 10 km range. This is important to discriminate between a constant with height profile and a large increase or decrease of the mole fraction below 5 km as suggested by some in situ measurements [Ignatiev et al., 1997; Donahue and Hodges, 1992]. The water mole fraction we derived is somewhat larger but still consistent, within error bars, with that determined at higher altitudes from the 2.3-mm (30 – 45 km) and 1.74-mm (15 – 25 km) nightside emission, which suggests that the water vapor mole fraction is constant below the clouds between 0 and 40 km, as predicted by chemical models [Krasnopolsky, 2007]. [30] We finally searched for spatial variations of the H2O abundance in various VIRTIS sequences covering altogether a large latitude range (80°S – 25°N). To do so, we used the ratio of the intensity in the two wings of the 1.18-mm window as a proxy. This search was negative and we derived stringent upper limits for any variation of the H2O mole fraction with latitude: ±1% for the range 60°S –25°N and ±2% if we consider the broader range 80°S – 25°N. These constraints are much stronger than those derived by Drossart et al. [1993] (±20%) from Galileo/NIMS spectra over a more limited spatial coverage. The horizontal uniformity of the H2O abundance is consistent with an abundance profile constant with altitude and with the expected lack of significant sources and sinks below 40 km. In fact, our results likely preclude any strong vertical gradient of the H2O concentration in the lowest scale height as vertical motions predicted by GCM calculations would then produce
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detectable horizontal variations as is the case for CO and OCS near 35 km [Marcq et al., 2006, 2008; Tsang et al., 2008]. [31] Acknowledgments. B.B., E.M., and P.D. acknowledge support from Centre National d’E´tudes Spatiales (CNES). G.P. was supported by Agenzia Spaziale Italiana (ASI). We gratefully thank the VIRTIS technical team for their continuous support.
References Be´zard, B., and C. de Bergh (2007), Composition of the atmosphere of Venus below the clouds, J. Geophys. Res., 112, E04S07, doi:10.1029/ 2006JE002794. Be´zard, B., C. de Bergh, D. Crisp, and J.-P. Maillard (1990), The deep atmosphere of Venus revealed by high-resolution nightside spectra, Nature, 345, 508 – 511, doi:10.1038/345508a0. Carlson, R. W., L. W. Kamp, K. H. Baines, J. B. Pollack, D. H. Grinspoon, T. Encrenaz, P. Drossart, and F. W. Taylor (1993), Variations in Venus cloud particle properties: A new view of Venus’s cloud morphology as observed by the Galileo Near-Infrared Mapping Spectrometer, Planet. Space Sci., 41, 477 – 485. Crisp, D. (1986), Radiative forcing of the Venus mesosphere: I. Solar fluxes and heating rates, Icarus, 67, 484 – 514, doi:10.1016/0019-1035(86)90126-0. Crisp, D., D. A. Allen, D. H. Grinspoon, and J. B. Pollack (1991), The dark side of Venus: Near-infrared images and spectra from the AngloAustralian Observatory, Science, 253, 1263 – 1266, doi:10.1126/science. 11538493. de Bergh, C., B. Be´zard, D. Crisp, J.-P. Maillard, T. Owen, J. Pollack, and D. Grinspoon (1995), Water in the deep atmosphere of Venus from highresolution spectra of the night side, Adv. Space Res., 15(4), 79 – 88, doi:10.1016/0273-1177(94)00067-B. Donahue, T. M., and R. R. Hodges Jr. (1992), Past and present water budget of Venus, J. Geophys. Res., 97, 6083 – 6091. Drossart, P., et al. (1993), Search for spatial variations of the H2O abundance in the lower atmosphere of Venus from NIMS-Galileo, Planet. Space Sci., 41, 495 – 504. Ignatiev, N. I., V. I. Moroz, B. E. Moschkin, A. P. Ekonomov, V. I. Gnedykh, A. V. Grigoriev, and I. V. Khatuntsev (1997), Water vapour in the lower atmosphere of Venus: A new analysis of optical spectra measured by entry probes, Planet. Space Sci., 45, 427 – 438, doi:10.1016/S00320633(96)00143-2. Jacquinet-Husson, N., et al. (1999), The 1997 spectroscopic GEISA databank, J. Quant. Spectrosc. Radiat. Transfer, 62, 205 – 254, doi:10.1016/ S0022-4073(98)00111-3. Krasnopolsky, V. A. (2007), Chemical kinetic model for the lower atmosphere of Venus, Icarus, 191, 25 – 37, doi:10.1016/j.icarus.2007.04.028. Marcq, E., T. Encrenaz, B. Be´zard, and M. Birlan (2006), Remote sensing of Venus’ lower atmosphere from ground-based IR spectroscopy: Latitudinal and vertical distribution of minor species, Planet. Space Sci., 54, 1360 – 1370, doi:10.1016/j.pss.2006.04.024. Marcq, E., B. Be´zard, P. Drossart, and G. Piccioni (2008), A latitudinal survey of CO, OCS, H2O, and SO2 in the lower atmosphere of Venus: Spectroscopic studies using VIRTIS-H, J. Geophys. Res., 113, E00B07, doi:10.1029/2008JE003074. Meadows, V. S., and D. Crisp (1996), Ground-based near-infrared observations of the Venus nightside: The thermal structure and water abundance near the surface, J. Geophys. Res., 101, 4595 – 4622, doi:10.1029/ 95JE03567. Moroz, V. I. (2002), Estimates of visibility of the surface of Venus from descent probes and balloons, Planet. Space Sci., 50, 287 – 297, doi:10.1016/S0032-0633(01)00128-3. Moroz, V. I., and L. V. Zasova (1997), VIRA-2: A review of inputs for updating the Venus International Reference Atmosphere, Adv. Space Res., 19, 1191 – 1201, doi:10.1016/S0273-1177(97)00270-6. Piccioni, G., et al. (2007), VIRTIS: The Visible and Infrared Thermal Imaging Spectrometer, Eur. Space Agency Spec. Publ., ESA SP 1295, 1 – 27. Pollack, J. B., et al. (1993), Near-infrared light from Venus’ nightside: A spectroscopic analysis, Icarus, 103, 1 – 42, doi:10.1006/icar.1993.1055. Seiff, A. (1983), Thermal structure of the atmosphere of Venus, in Venus, edited by D. M. Hunten et al., pp. 215 – 279, Univ. of Ariz. Press, Tucson. Stamnes, K., S. C. Tsay, W. Wiscombe, and K. Jayaweera (1988), Numerically stable algorithm for discrete-ordinate-method radiative transfer in multiple scattering and emitting layered media, Appl. Opt., 27, 2502 – 2509, doi:10.1364/AO.27.002502. Taylor, F. W., D. Crisp, and B. Be´zard (1997), Near-infrared sounding of the lower atmosphere of Venus, in Venus II: Geology, Geophysics, Atmo-
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sphere, and Solar Wind Environment, edited by S. W. Bougher et al., pp. 325 – 351, Univ. of Ariz. Press, Tucson. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017. Tonkov, M. V., N. N. Filippov, V. V. Bertsev, J. P. Bouanich, N. Van-Thanh, C. Brodbeck, J. M. Hartmann, C. Boulet, F. Thibault, and R. Le Doucen (1996), Measurements and empirical modeling of pure CO2 absorption in the 2.3-mm region at room temperature: Far wings, allowed and collisioninduced bands, Appl. Opt., 35, 4863 – 4870, doi:10.1364/AO.35.004863.
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Tsang, C. C. C., et al. (2008), Tropospheric carbon monoxide concentrations and variability on Venus from Venus Express/VIRTIS-M observations, J. Geophys. Res., 113, E00B08, doi:10.1029/2008JE003089.
B. Be´zard, P. Drossart, and E. Marcq, LESIA, Observatoire de Paris, 5 place Jules Janssen, F-92190, Meudon, France. (Bruno.Bezard@ obspm.fr) R. W. Carlson, Jet Propulsion Laboratory, MS 183-601, 4800 Oak Grove Drive, Pasadena, CA 91109, USA. G. Piccioni, INAF, IASF, via del Fosso del Cavaliere 100, I-00133 Rome, Italy. C. C. C. Tsang, Atmospheric, Oceanic and Planetary Physics, Department of Physics, University of Oxford, Clarendon Laboratory, Parks Road, Oxford, OX1 3PU, UK.
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Venus cloud top winds from tracking UV features in Venus Monitoring Camera images R. Moissl,1 I. Khatuntsev,1,2 S. S. Limaye,3 D. V. Titov,1,2 W. J. Markiewicz,1 N. I. Ignatiev,2 T. Roatsch,4 K.-D. Matz,4 R. Jaumann,4 M. Almeida,5 G. Portyankina,1 T. Behnke,4 and S. F. Hviid1 Received 15 February 2008; revised 9 October 2008; accepted 10 November 2008; published 17 March 2009.
[1] To date dynamical observations of the Venus clouds have delivered mainly either only
short-term or long-term averaged results. With the Venus Monitoring Camera (VMC) it finally became possible to investigate the global dynamics with a relatively high resolution in space and time on a long-term basis. Our findings from manual cloud feature wind tracking in VMC UV image sequences so far show that the details of the mesospheric dynamics of Venus appear to be highly variable. Although the general rotation of the atmosphere remained relatively stable since Mariner 10, more than 30 years ago, by now, there are indications of short-term variations in the general circulation pattern of the Venus atmosphere at cloud top level. In some cases, significant variations in the zonal wind properties occur on a timescale of days. In other cases, we see rather stable conditions over one atmospheric revolution, or longer, at cloud top level. It remains an interesting question whether the irregularly observed midlatitude jets are indeed variable or simply become shielded from view by higher H2SO4 haze layers for varying time intervals. Winds at latitudes higher than 60°S are still difficult to obtain track because of low contrast and scarcity of features but increasing data is being collected. Over all, it was possible to extend latitudinal coverage of the cloud top winds with VMC observations. Thermal tides seem to be present in the data, but final confirmation still depends on synthesis of Visible and Infrared Thermal Imaging Spectrometer and VMC observations on night and dayside. Although poorly resolved, meridional wind speed measurements agree mainly with previous observations and with the presence of a Hadley cell spanning between equatorial region and about 45°S latitude. Citation: Moissl, R., et al. (2009), Venus cloud top winds from tracking UV features in Venus Monitoring Camera images, J. Geophys. Res., 114, E00B31, doi:10.1029/2008JE003117.
1. Introduction [2] The cloud level super rotation of the atmosphere, first determined from ground based images [Boyer and Guerin, 1969] has been measured from Mariner 10 [Limaye and Suomi, 1981], Pioneer Venus and Galileo missions [Rossow et al., 1990; Belton et al., 1991; Toigo et al., 1994; Limaye, 2007; Peralta et al., 2007]. [3] During its flyby at Venus in 1974 Mariner 10 collected image data with resolutions down to 30 km/pixel during the 8 days around its closest approach to the planet. The most 1 Max-Planck-Institut fuer Sonnensystemforschung, Katlenburg-Lindau, Germany. 2 Space Research Institute, Moscow, Russia. 3 Space Science and Engineering Center, University of WisconsinMadison, Madison, Wisconsin, USA. 4 Institut fur Planetenforschung, Deutsches Zentrum fur Luft- und Raumfahrt, Berlin, Germany. 5 European Space Astronomy Centre, European Space Agency, Madrid, Spain.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003117$09.00
extensive data set so far was obtained by the Pioneer Venus OCPP instrument between 1979 and 1986. The Galileo probe, which performed a flyby in 1990, acquired data during the 16 h around closest approach to Venus. The UV filter was centered at a wavelength of 400 nm. [4] The super rotation refers to the fact that bulk of the atmosphere of Venus rotates faster than the underlying solid planet from the surface to an altitude of 80 km, reaching speeds above the visible cloud top at rates 50 times faster than the planet. How this super rotation is maintained has been a puzzle since its discovery. The combination of the strong zonal flow and the weaker meridional flow explains the spiraling streaks seen in the ultraviolet images [Suomi, 1975; Smith and Gierasch, 1996] and is likely responsible for, or an artifact of the hemispheric vortex organization of the circulation centered over each pole of Venus [Limaye and Suomi, 1981; Limaye, 2007]. [5] The inferred global circulation is that the atmosphere is organized vertically in at least one Hadley circulation cell, wherein the solar heating in low latitudes causes rising motion and near the cloud level atmosphere flows toward the pole where radiative cooling leads to sinking to com-
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plete the return flow (i.e., thermally direct circulation). The indirect evidence of this cell extending to the limiting altitude of the observations (40 km) was determined from the detection of the solenoidal circulation from the temperature profiles determined from the radio occultation technique from the Pioneer Venus orbiter through the determination of the angle between the pressure and density surfaces [Limaye, 1985]. Winds inferred from tracking entry probes are inconclusive about whether there may be multiple cells between the surface and the 40 km level below which the thermal structure data from radio occultations are not available, but have been postulated [Schubert, 1983]. [6] The previous studies of the cloud motions have revealed aspects of the cloud level circulation such as Kelvin and gravity waves [Del Genio and Rossow, 1990; Smith and Gierasch, 1996] and presence of thermal tides [Limaye et al., 1982; Rossow et al., 1990; Limaye, 2007]. Investigations on these phenomena are part of the continuing effort to understand the processes that maintain the super rotation of the atmosphere [Schubert, 1983; Schubert et al., 2007] These include observations of the circulation below the clouds from entry probes or balloons and, more recently, on the night side from tracking of features in near infrared observations of Venus from Galileo [Carlson et al., 1991; Sa´nchez-Lavega et al., 2008] and Venus Express as well as from telescopes on Earth [Crisp et al., 1991; Limaye et al., 1988]. [7] Schubert [1983] discusses the different suggestions for the maintenance of the super rotation of the atmosphere, which include processes such as equatorward transport of angular momentum by eddies and angular momentum transport through solar thermal tides. Thus much of the effort continues to be directed at establishing the roles of eddies and solar tides in the transport of angular momentum. Limaye [2007] discussed the challenges posed by the lack of sufficient and complete observations due to the fact that until recently the thermal tides in the winds could be detected only from the cloud motions measured on the day side, which likely introduced a bias in the estimates of meridional transport of angular momentum by the mean and eddy circulation if the night side wind distributions were strong and different enough to cause the true zonal average to be substantially different from the day side average. The absence of night side information also impacts the inferences about the Kelvin waves [Del Genio and Rossow, 1990] to some degree and in the understanding of true temporal variations in the cloud level circulation on Venus. [8] One of the main goals of the Venus Express mission is to provide a global and systematic study of the atmospheric circulation [Svedhem et al., 2007; Titov et al., 2006]. We use sequences of UV images taken by Venus Monitoring Camera [Markiewicz et al., 2007a, 2007b] to measure the wind speeds at the cloud tops by tracking apparent motions of the cloud features. In comparison to the earlier investigations these observations provide extended and systematic latitude and local time coverage including high latitudes in nadir viewing geometry, significantly better temporal resolution and long-term coverage, and have higher spatial resolution. The preliminary results of the UV digital cloud tracking were presented by Markiewicz et al. [2007b]. Here we describe in detail the Venus Monitoring Camera (VMC) imaging sequences, the cloud tracking techniques, and the
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results of cloud tracking in VMC images from two slightly different manual tracking techniques during the primary mission of Venus Express. The digital tracking results will be briefly addressed in this work and presented in more detail at a later date pending further analysis.
2. Observations and Data Set [9] The Venus Monitoring Camera is a wide angle camera with a field of view of 17°, imaging Venus simultaneously in four narrow-band filter channels. All four channels are sharing one CCD detector, each one using one quadrant of the CCD. Two narrow band filters are in the near-IR (965 nm and 1010 nm), one in the visible (513 nm) and one in UV (365 nm). The UV channel is centered at the spectral signature of the unknown UV absorber. We use this filter to observe and track the UV cloud markings on the dayside, which are detected at around 70 km of altitude, where the optical depth in the UVapproaches unity [Tomasko et al., 1985; N. I. Ignatiev et al., Altimetry of the Venus cloud tops from the Venus Express observations, submitted to Journal of Geophysical Research, 2008]. A more detailed description of the VMC instrument, its properties and science goals has been presented in the work of Markiewicz et al. [2007a; 2007b]. [10] Because of the highly elliptical polar orbit of the Venus Express spacecraft the distance from the planet varies from 250 km at pericenter to 66.000 km in apocenter in the course of one 24 h orbit. Therefore the VMC angular resolution of 0.74 mrad per pixel translates into spatial resolution of approximately 0.2 km/pixel at pericenter and 50 km/pixel around apocenter. Consequently, VMC data yields high-resolution close-up images of the Northern Hemisphere as well as a global view of Venus from the South Pole. [11] In October 2007 Venus Express completed its primary mission [Svedhem et al., 2007; Titov et al., 2006]. In the course of 510 revolutions about the planet, the Venus Monitoring Camera acquired a total inventory of roughly 67,000 images with 31,800 in the UV among them. The data and results presented in this paper are based on the images obtained at 60,000 – 26,000 km distance from Venus, corresponding to a pixel size of 20– 50 km. Typical time intervals between image pairs used for cloud tracking are 40 min that corresponds to 5 – 10 pixel displacement of cloud features between the images. This selection is a compromise between the error in velocity measurements and temporal resolution. The planet coverage, regarding tracked UV markings, ranges from about 10°N to 80°S in latitude and 8 – 16 h local time. The contrast in VMC UV images varies between 5% and 30%, with typical values of 15– 30% for the tracked features. [12] Because of some minor damage to the detector due to solar irradiation, all images are subject to an in-flight flat fielding procedure before evaluation. In order to eliminate all artifacts in the VMC data, a series of images is acquired at high latitudes where the observed contrast is at a minimum. The images of these sequences are then averaged to achieve a uniformly gray background on which all persistent artifacts stand out prominently. All images in the corresponding orbit are then divided by the (normalized)
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Figure 1. Coverage by VMC cloud-tracking measurements per 5° latitude 0.5 h local time bin. flat field. Because of a temporal evolution of some artifact patterns minor residuals may remain. [13] The Venus Express orbit can be divided roughly into three parts: apocentric (66,000 – 50,000 km distance from the planet), ascending branch (50,000 – 10,000 km) and pericentric (10,000 – 250 km). In these three orbital segments VMC images are acquired in order to study various types of features and observe different aspects of the cloud top dynamics. In apocentric images Venus is seen almost directly from beneath the South Pole, so the polar region is imaged without any foreshortening (unlike previous missions), but the low latitudes are seen in oblique view. This part of the VMC data set is used to study near-polar global dynamics and cloud morphology. In the ascending branch VMC observes middle and low latitudes in nadir geometry. Local time coverage depends on the sub spacecraft longitude that slowly changes from one orbit to another (1.5 deg or 6 min per day). The ascending branch images are predominantly used to track winds from medium and small-scale features, ranging from about thousand down to several hundred kilometers in diameter. At close approach cloud features on scales down to few tens of kilometers in size can be identified in the images. However, at distances closer than 10,000 km wind tracking becomes increasingly more difficult as the rapid motion of the spacecraft decreases the overlap between images and features can be kept in the field of view only for short times. [14] Figure 1 shows the coverage of the latitude-localtime field by the VMC cloud tracking. The number of vectors peaks near downstream (to the west) of the sub solar region where a large number of discrete features of convective nature is regularly observed and drops toward higher latitudes as well as the morning and evening terminator. The main limiting factor for coverage in local time in the dusk
and dawn regions are the steep brightness gradients which, even in high-pass-filtered images, dampen contrast of the UV markings. At polar latitudes the morphology of the clouds is different altogether, presumably due to a decrease in or even absence of convective activity and presence of submicron haze [Kawabata et al., 1980]. This results in lack of discrete features, very low contrasts and drawn-out diffuse feature boundaries. In total, about 20,000 vectors have been extracted from 450 image pairs. [15] During the primary mission the descending branch of the Venus Express orbit was reserved for data transmission to Earth. In conjunction with the relative change in local solar time of the orbit plane, this lead to separation of the VMC wind tracking data into three periods in which it was possible to observe the dayside of Venus in ascending branch from a sufficient distance. These three periods are separated by gaps of almost 140 orbits in which VMC was pointed either to the night side or too close to the terminator to allow for tracking sequences to be taken in the required orbit segment. The first period spans from orbit 29 to 72 (containing tracking sequences from orbits 29, 30, 31, 34, 38, 46, 51, 56, 60, 61, 72, May– June 2006), the second period from orbit 208 to 295 (containing 208, 210, 230, 246, 250, 257, 258, 260, 263, 265– 267, 279, 281– 284, 295, November 2006 to January 2007), the third period from orbit 439 to 530 (containing 439, 440, 442, 453, 460– 463, 469, 471, 530, July – September 2007). Table 1 is listing the according dates of each orbit in the data set, together with the approximate coverage of the wind tracking in local time (LTR) and latitude (LAR), as well as the local time of the subspacecraft point at ascending node (LTAN). [16] Usually the individual wind tracking sequences are separated from each other by several days. But also sets of up to 4 consecutive orbits are present in the data set. These
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Table 1. List of All Orbits in the Data Seta Orbit
Date
LTR (h)
LAR (°)
LTAN (h)
29 30 31 34 38 46 51 56 60 61 72 208 210 230 246 250 257 258 260 263 265 266 267 279 281 282 283 284 295 439 440 442 453 460 461 462 463 469 471 530
19 May 2006 20 May 2006 21 May 2006 24 May 2006 28 May 2006 5 Jun. 2006 10 Jun. 2006 15 Jun. 2006 19 Jun. 2006 20 Jun. 2006 1 Jul. 2006 15 Nov. 2006 17 Nov. 2006 7 Dec. 2006 23 Dec. 2006 27 Dec. 2006 3 Jan. 2007 4 Jan. 2007 6 Jan. 2007 9 Jan. 2007 11 Jan. 2007 12 Jan. 2007 13 Jan. 2007 25 Jan. 2007 27 Jan. 2007 28 Jan. 2007 29 Jan. 2007 30 Jan. 2007 10 Feb. 2007 4 Jul. 2007 5 Jul. 2007 7 Jul. 2007 18 Jul. 2007 25 Jul. 2007 26 Jul. 2007 27 Jul. 2007 28 Jul. 2007 3 Aug. 2007 4 Aug. 2007 2 Oct. 2007
10.0 – 15.1 08.8 – 16.5 09.3 – 16.7 08.2 – 16.2 08.9 – 15.5 08.1 – 15.8 09.3 – 15.6 10.1 – 16.7 09.4 – 16.4 11.5 – 16.3 09.7 – 17.4 07.8 – 15.4 07.8 – 14.6 08.0 – 15.3 08.0 – 16.3 07.8 – 16.0 08.6 – 16.3 08.8 – 15.8 08.1 – 16.4 08.3 – 16.6 09.9 – 16.3 08.6 – 16.3 10.6 – 16.4 07.9 – 17.2 09.5 – 17.2 08.5 – 17.2 09.0 – 17.0 10.3 – 16.8 09.6 – 17.0 08.2 – 15.3 08.2 – 16.1 07.6 – 15.2 07.1 – 16.7 07.3 – 15.9 08.6 – 15.8 08.0 – 16.7 07.4 – 16.7 07.6 – 16.3 07.8 – 16.7 08.9 – 17.2
78 to 15 78 – 01 76 – 13 79 to 02 65 – 05 81 – 05 79 to 02 69 – 00 73 – 07 67 – 12 74 – 02 73 – 04 75 – 05 73 – 04 71 – 08 78 – 09 67 – 11 75 – 07 76 – 05 77 – 11 75 – 10 72 – 15 73 – 06 82 – 08 73 – 11 84 – 13 77 to 06 76 – 07 80 – 01 78 to 02 80 – 01 79 – 00 82 – 01 74 – 04 72 to 02 73 – 06 74 – 01 74 – 08 75 – 03 74 to 03
13.6 13.7 13.8 14.1 14.6 15.4 15.9 16.5 16.9 17.0 18.2 8.8 9.0 11.2 12.8 13.3 14.0 14.1 14.3 14.6 14.8 14.9 15.0 16.3 16.5 16.6 16.7 16.8 18.0 9.5 9.6 9.8 11.0 11.7 11.8 11.9 12.0 12.7 12.9 19.1
a
The LTR column shows the approximate coverage of the wind tracking in local time (decimal representation). The LAR column gives approximate latitudinal coverage (negative values denote southern latitudes). LTAN refers to the subspacecraft local time at ascending node.
small irregularities in coverage were compensated by the fact that the field of view included broad range of local times in distant imaging.
3. Wind Tracking Procedures 3.1. Description of the Employed Tracking Methods [17] We used both digital (automatic) and visual (manual) tracking of the UV cloud markings to derive the wind speeds. Three independent experimenters used imaging sequences from 39 orbits and applied two different visual tracking techniques and automatic wind tracking software. The first step in all cloud tracking methods was to establish navigation for the used images. We used the postprocessed SPICE data (see: http://naif.jpl.nasa.gov/naif/) to calculate VMC boresight vectors and their intersection with the cloud layer, assigning planetary coordinates to each image pixel. The cloud top height was set to a fixed value of 70 km above the planetary surface for navigation purposes. We expect the actual cloud top height to vary by about ±5 km
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around this level, which results in negligible uncertainties in coordinate determination. [18] For historical reasons and software availability we use two different procedures for visual wind speed measurements. The ‘‘sequential method’’ traces back to the wind speed measurements from previous missions, especially from Pioneer Venus OCPP observations [Limaye, 1988; Limaye et al., 1982, 1988]. As done for the OCPP wind speed measurements a loop of an entire image sequence from one VEX orbit was displayed in either equiangular (rectilinear, 0.2°/pixel) or polar projection onto a fixed coordinate grid (Figure 2 (right)). In this method motion of a cloud marking is observed in the way of a short movie clip, giving the experimenter an overview of the path of a selected cloud feature as well as the global pattern of motion. After familiarizing with the general flow pattern the experimenter puts small markers on distinctive cloud features in one image from the given orbit. Then the loop is continued, switching to a later image from the same sequence. There, another group of markers is put on the same cloud features. These steps are then repeated until the selected cloud features have changed significantly from their initial shapes and sizes and cannot be identified anymore. To enhance visibility of the cloud features a contrast enhancing high pass filter had been applied to the images after care was taken that the filtering did not alter any other image properties. This technique was used by two observers for about 40% (8000) cloud vectors. [19] The ‘‘paired method’’ is a straightforward approach in which two different images from the same sequence are displayed side by side on a monitor and the cloud features in both are observed and compared simultaneously (Figure 2 (left and middle)). Also in this method, markers are set in order to track the corresponding points of cloud features in the compared image pairs. One significant difference is that here the images did not undergo any filtering, contrast enhancement or projection. Thus, basically unaltered VMC images were used for tracking with this method. The images were only slightly resized in order to achieve an approximately constant display size of the planet in each image. The paired method was used by one observer for about 60% (12,000) of cloud vectors. [20] Although the basic principle, visual correlation between two images, is the same in both approaches, there are subtle but important differences. In the first case the observer tracks clouds by blink comparison between two images, trying to follow displacement of the tracked features on the fixed coordinate grid. In the second case the tracked feature is seen in both images simultaneously, thus emphasizing visual pattern recognition. [21] In both methods the experimenters tracked a great variety of feature types. Basically all features that did not span on a globally significant scale were used to track the cloud movements. We did not differentiate between ‘‘bright’’ and ‘‘dark’’ features, tracking both alike. The main criteria for feature selection were as follows: (1) Significant, well discernible contrast between the feature and its immediate surroundings. No fixed criterion was applied but the vast majority of tracked features had contrasts in excess of 15 % with respect to their surrounding. (2) Features size of 100– 300 km that allows for reliable identification of the feature center (only in equatorial regions). (3) Prominent
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Figure 2. Typical VMC UV images used for tracking purposes. (left and middle) Pair of original VMC images used in the paired method. (right) Image in rectilinear projection used in sequential method. shapes like protrusions or clearly edged parts of boundaries for larger features. (4) Minimum time intervals between images were chosen to allow for resolving wind speed differences of about 10 – 20 m/s (depending on image resolution). (5) No features from the limb regions were used because they are subject to foreshortening. [22] The wind speeds in both methods were derived in similar ways. First the navigation data was linked to the images (for both, the projected and unprojected images). Then latitude and longitude of the respective marked pixels were extracted for all measurements of the sequence. From these values and the time intervals between images the zonal and meridional cloud speeds were calculated as U ðl; t; 8Þ ¼ DlRC =Dt cosð8Þ; V ð8; tÞ ¼ D8RC =Dt;
where U and V are zonal and meridional wind speed components, l and 8 are longitude and latitude, RC is the distance from the planetary center to 70 km above the surface (RC = 6121.8 km), and Dt is the time between images. [23] For the digital tracking in VMC images we used the same algorithm as applied by Rossow et al. [1990]. The analysis was performed systematically in pairs of images with at least one hour time difference using cross correlation as a metric with a ‘‘cloud’’ target of 10 10 pixels in the first image (2° 2° latitude by longitude). The displacement between the target location in the first image and the location of the maximum cross correlation coefficient in the second image between the search and the target was taken as the best estimate of the movement of the selected target over the interval between both images. The displacement over the two images leads to the determination of the zonal and meridional components of the cloud level flow. The search window in pixels was set to a window equivalent to a maximum flow of ±100 m/s in the meridional and 200 m/s in the zonal direction, accounting for the time interval between the images. Here the digital tracking results on 205,000 vectors obtained from images acquired on nine orbits (30, 31, 34, 38, 46, 51, 56, 60, and 61) are presented in Figure 3 to show the generally good agreement between the results from visual and digital tracking. Preliminary results from orbits 262– 267 were presented by Markiewicz et al. [2007b].
[24] Figure 2 shows examples of images used in the two different visual tracking methods. After the tracking sequences have been compiled from the orbital data set and markers have been set, the marker positions are determined and recorded. From the resulting set of vectors, latitudinal and/or longitudinal wind profiles and maps can be determined, by averaging all wind vectors inside a latitude bin. From the average density of retrieved vectors per orbit, we decided to use latitude bins of 5° and longitude bins of 7.5° (corresponding to 0.5 h of local solar time) in order to maintain good levels of statistics. 3.2. Comparison of Wind Tracking Measurements [25] As is obvious from Figure 2, the appearance of the cloud deck strongly differs between the original and the projected images. While the original images show all parts of the planet true to the perspective of the VMC, the rectilinear projection tends to distort the polar region quite drastically. Respectively, the polar projection distorts the equatorial region to a point where features get too distorted to be used for tracking. The basic properties of both methods will be addressed below. [26] Since Venus cloud features evolve over time, visual tracking relies on a somewhat subjective approach to identify a cloud feature in two or more images acquired at different times. In the case of VMC data, the orbital motion of the Venus Express spacecraft results in a slightly different ‘‘view’’ of a given cloud because of both its displacement due to ambient flow and due to the change in perspective. The paired method thus visually measures the combined displacement from perspective and ambient flow. And the part due to perspective change is accounted for by the image navigation implicitly. From the preserved perspective in the original images, the experimenter has good insight into which parts are seen heavily foreshortened, especially close to the limbs. This allows for a well founded possibility to select only the weakly or nonforeshortened features for tracking. On the other hand the apparent feature size and resolution change along the orbit. Therefore one has to be careful about reidentifying the same part of a feature in different images. This is especially the case in the lowcontrast areas at higher latitudes. There it is possible for faint feature boundaries to interfere with so-called ‘‘flat field
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Figure 3. Latitude profiles of zonal winds derived from the paired (red line) tracking, sequential (blue line) visual tracking, and digital (black line) tracking. Error bars represent standard deviation due to uncertainties of individual measurements.
remainders’’ which result from the flat field correction process applied to VMC images. [27] In the sequential tracking method the flat field remainders become distinctively visible after high pass filtering, because they represent high-frequency variations. So they can be easily identified and avoided, since their apparent movement is governed by the change of the VMC perspective on the planet. Thus they mimic a movement completely different from all real atmospheric flows. [28] Figure 3 compares the latitude profiles, averaged in 5° latitude bins, of zonal wind speeds derived from the agglomerate data sets of the sequential (red curve) and paired (blue curve) visual tracking methods and those from digital tracking (black curve). The sequential method shows slightly higher speeds on average, which might be an indication for minimal bias due to either method. The biggest differences occur at high latitudes, between 70 and 80°S. Although the two methods do not actually agree there very well within their respective measurement accuracies, this is more likely due to the difficulties in finding adequate features for tracking in these latitudes than to differences in the tracking methods. Therefore, these deviations seem to originate from observer bias. Regarding the above mentioned risk of interference of faint, streaky cloud features with flat field remainders, some influence of this interference might possibly also play a role in this. Also the digital and the visual tracking results from the VMC data are in very good agreement up to about 40° latitude, and show a difference of 15 or 20 ms1 between 40 and 60° latitude from the visual tracking. There are likely two causes for this (1) the morphology and low contrast at these latitudes makes pattern matching by digital cross correlation difficult and (2) a selection effect in that
visually only the targets that are moving fast are discernible, but in longitude regions where the flow may be slower, no targets are seen as is suggested by the longitudinal distribution of the visual vectors. The digital results south of 60°S are not very reliable yet and are therefore to be viewed as entirely preliminary at this point. [29] Further, the average standard variation in all latitude bins is approximately the same for both visual tracking methods. This is a clear indication that measurement accuracy is nearly the same for both methods. [30] In conclusion, the deviations between the two visual methods are well on the order of the expected measurement uncertainty for single measurements in the according latitude regions. Despite the obvious discrepancy in the 70– 80°S latitude range between the two profiles from the two visual tracking methods, we decided that merging also these results will improve the poor statistics in this region and minimize observer bias. From this we conclude that the results from both methods are in reasonable agreement to allow for regarding all measured wind vectors as one data set, regardless of the employed method. In consequence later on in the paper we do not differentiate the results by tracking method. 3.3. Error Sources and Systematic Uncertainties [31] A number of measurement properties are influencing the accuracy of the wind speed measurements. Main sources of uncertainty are image resolution, measurement accuracy and feature evolution. Additionally some bias from the different methods and/or observers might introduce systematic errors. [32] Velocity uncertainty of an individual measurement due to technical limitations, such as image resolution and
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marker position error in the image is easily assessed via first-order Taylor expansion as follows: dU ¼
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nounced features and up to three pixels for fuzzy lowcontrast features in high latitudes.
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ffi pffiffiffi ð 2Rc da cosð8Þ=DtÞ2 þ ðdaRC sinð8Þ=DtÞ2 þ ðadRc cosð8Þ=DtÞ2 qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi pffiffiffi dV ¼ ð 2Rc da=DtÞ2 þ ðadRc =DtÞ2 ;
where d values denominate total uncertainties for velocities (U,V), measured angular differences (a), and distance from Venus center to the cloud top (RC) and 8 is the latitude at measuring point. [33] For equatorial latitudes the first term, assuming a 1 pixel error (corresponding to approximately 0.2– 0.4° in latitude and longitude, depending on observing distance) and 40 min between images further away and 20 min when closing in, ranges from about 36 m/s at apocenter to roughly 10 m/s at 30,000 km distance from the planet, being 20 m/s on average. At higher latitudes the accuracy in latitudinal direction decreases further because of the convergence of latitude circles. This accuracy limit poses a great problem for deriving reliable meridional speeds, since they are supposed to be of the order of 10 m/s. Of course this source of uncertainty could be further suppressed through choice of image pairs with further increased time intervals. The early tracking sequences where too short to allow for reasonable statistics with image pairs separated by more than one hour, thus the chosen time intervals represent a trade-off between statistics and precision of the measurements. The second term, assuming 3 – 5 pixels between images and 10 km uncertainty in cloud deck height, remains well below 1 m/s in any case and can therefore be neglected. [34] Sources for uncertainty of an individual measurement could arise from the following: [35] 1. Pointing and mapping inaccuracies are considered to be negligible since both have proven to be extremely accurate and stable. A very conservative estimate would be an inaccuracy amounting to one pixel. [36] 2. Pixelation and noise are assumed to amount to an error of one pixel if physical feature boundaries lie close to the edge of a pixel. Depending on flat field quality, which is varying from orbit to orbit since images of nearly featureless cloud areas are needed, the uncertainty by noise can reach up to two pixels. [37] 3. Morphology and morphological evolution also influence the measurement accuracy. Since cloud features seldom have clear boundaries, there is always some possible (and variable) inaccuracy in determining the center or exact boundary line of a tracked feature. Furthermore, evolution of cloud features is observed, but because of time intervals between images being always smaller than 4 h, no significant inaccuracy should arise from them. This has been discussed already for the Pioneer Venus OCPP data set [Rossow et al., 1990]. These issues are especially valid for higher latitudes, since the low contrasts there make identification of the feature boundaries more difficult. It is difficult to quantify the influence of these morphological uncertainties, since they vary significantly with latitude and feature type. As a rule of thumb, the corresponding uncertainty varies between one pixel for small, sharply pro-
[38] 4. Systematic and random errors from measurement methods are likely to have some effect, as mentioned above. To minimize the effects of erroneous wind speeds from faulty measurements, we decided to reject all vectors outside the interval of 160– 0 m/s. The constraints for meridional wind speeds were 60– +60 m/s. These selection criteria lead to rejection of about 1.4% of all measured vectors. [39] As in previous works about cloud top winds inferred from the cloud feature tracking, we find that the variability in the derived and the averaged wind speed profiles are larger than the expected uncertainties from individual measurements. Therefore we chose to apply the standard deviation of wind speeds in one latitude bin as a measure of uncertainty for all wind speed profiles.
4. Results of the Wind Tracking 4.1. Zonal Wind Profiles [40] Figure 4 shows the averaged latitude profiles of zonal wind speed for the aggregate visual and digital tracking data sets, binned in latitude bins of 5° width. Also shown in Figure 4 is the profile obtained from the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) 380 nm channel data [Sa´nchez-Lavega et al., 2008] and the corresponding rotation period derived from the digital and visual profiles. UV winds derived from both experiments are in good agreement although VIRTIS tends to show systematically stronger winds by 10 m/s. At low latitudes, zonal wind speed is 85 – 90 m/s and almost constant with latitude. The latitude profile shows a gradual increase to 100 m/s, peaked at 45°S, indicating the presence of a weak midlatitude jet which is also seen in the inferences from the VIRTIS thermal wind analysis [Piccialli et al., 2008]. South of this latitude the wind speeds steadily decrease toward the pole. Although the profile is close to a solid body rotation curve at first glance, a closer look on the rotation period in Figure 4b reveals significant differences, where solid body rotation profiles would show up as horizontal lines. The kink at 10°N in the zonal wind profile is very likely an artifact, due to the rather poor sampling as it is at the very edge of VMC coverage of Venus in the tracking image sequences. [41] The error bars represent standard deviation of the zonal wind speed in each latitude bin. The deviation increases toward higher latitudes, mostly due to the changes in morphology and difficulties in finding well defined UV markings there, as discussed above. The standard deviations exceed the uncertainties of individual measurements orbit (compare to Figure 3), indicating a likely orbit to orbit variability of the latitudinal profiles of the zonal wind component. Also variability with local time is likely to be present because of influences from planetary-scale waves [Del Genio and Rossow, 1990].
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Figure 4. (a) Average zonal wind speed profiles from VMC visual (solid line) and digital (dotted line) in comparison with data from VIRTIS (dashed line) [Sa´nchez-Lavega et al., 2008]. VMC and VIRTIS data is binned by 5° and 2°, respectively. The error bars represent standard deviation of the whole data set, including both measurement inaccuracies and orbit-by-orbit variations. (b) Corresponding rotation period profile from VMC wind tracking data (solid line is visual and dashed line is digital). [42] Figure 5 shows a selection of different single orbit profiles and the average profiles for each of the three observation periods. Standard deviations per latitude bin divided by the respective square root of the number of measurements are shown in the bottom Figures 5a and 5b, giving a measure of the standard error for the average value per bin. The mean zonal speeds can change by as much as 30 m/s from orbit to orbit and show significant variability in strength, position and even presence of the mid latitude wind speed maximum. Since these differences are well outside the standard error, they are most likely representing true variations of the zonal flow. These rather strong shortterm variations in the wind speed profiles leads to slightly different characteristics in the average profiles for each of the three observation periods. In periods 1 and 3, pro-
nounced wind speed maxima have been observed in the mid latitude regions, which is not the case for period 2. 4.2. Meridional Wind Component [43] Since the measured tracking sequences do not allow for individual measurements with accuracies better than ±10 m/s, all results on the meridional winds have to be considered preliminary. The standard deviation in each 5° latitude bin amounts to ±13 m/s. These are about as large as the uncertainties expected to arise from the chosen combination of image resolutions and time intervals. This indicates that measurement inaccuracies are likely larger than the real variability in meridional wind speeds. Better accuracy meridional component profiles are possible from extended mission data and will be published in future.
Figure 5. Latitude profiles of zonal wind for (a) individual orbits and (b) means for three tracking periods. The lines in the bottom of the images show average standard deviation divided by the square root of the number of measurements of the compared profiles for each latitude bin. 8 of 13
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Figure 6. Average meridional wind profile. Error bars represent RMS deviation of measurements in one bin. [44] As one can see from Figure 6 the average meridional wind component ranges from 0 to 10 m/s, which is roughly consistent with previous results. Meridional speeds are slightly increasing from about 5 m/s at the equator to peak values of 10 m/s at mid latitudes between 40 and 55°S and then decreasing again toward the pole. The observed profile is in agreement with the assumption of a Hadley cell circulation between the equator and the mid latitudes [Schubert, 1983; Gierasch et al., 1997]. [45] Future tracking sequences acquired at shorter distances from Venus will help to improve on accuracy of the
meridional wind measurements by increasing image resolution which until now is not sufficient to satisfactory resolve wind speeds on the order of 0 – 10 m/s. We hope to be able to investigate temporal variability once higherresolution tracking sequences have been evaluated. 4.3. Dependence of the Zonal Component on Local Solar Time [46] Figure 7 shows the average zonal wind field as a function of latitude and local time after binning the wind speed vectors in ‘‘boxes’’ of 0.5 h local solar
Figure 7. Local solar time versus latitude field of the zonal wind in the Southern Hemisphere. 9 of 13
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Figure 8. Local time profile of average zonal wind speeds between 0 and 20°S. VMC measurements in comparison with best fit for diurnal and semidiurnal tidal model (solid line) and diurnal model only (dashed line). time 5° ltitude each. The zonal wind field has a pronounced minimum at 11.5 h slightly upstream of the subsolar point and accelerates in the afternoon in low latitudes. In the middle latitudes velocities are higher on average over the whole observed local time range and show less local time variability. In higher latitudes an increase in speed with local time is apparent down to 60°S. Toward the pole zonal speeds are decreasing rapidly. [47] The data displayed in the Figure 7 has been divided into three latitude bands (0 – 20°S, 20 –40°S, and 40– 60°S) in order to evaluate tidal components in the zonal wind field
in the different latitude regions (Figures 8, 9, and 10). In each of the latitude bands we compared the local time profiles to a tidal model that includes both diurnal and semidiurnal components U ðlÞ ¼ U0 þ U1 sinðl þ F1 Þ þ U2 sinð2l þ F2 Þ;
where U0 is the zonal mean value, l is the longitude, U1 and U2 are the diurnal and semidiurnal velocity amplitudes, and F1 and F2 are the diurnal and semidiurnal phase angles.
Figure 9. Local time profile of average zonal wind speeds between 20 and 40°S. VMC measurements in comparison with best fit for diurnal and semidiurnal tidal model (solid line) and diurnal model only (dashed line). 10 of 13
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Figure 10. Local time profile of average zonal wind speeds between 40 and 60°S. The dashed line indicates the general trend with local time.
[48] In Figure 8, the local-time-dependent zonal wind component in the latitude strip between 0 and 20°S is being compared to fits from the diurnal-only part of the model and both diurnal and semidiurnal. Both fit curves show considerable agreement with the observed wind speeds, with higher agreement when both tidal wave modes are employed. Figure 9 shows the same comparison for the 20– 40°S latitude strip. Here the data is in about equal agreement with both model curves. For the high-latitude region no conclusive correlation for diurnal and semidiurnal wave models could be obtained. The data shows, however, an overall trend toward increased wind speeds in the afternoon hours. [49] In summary the observed local time variations of the zonal wind component is indicating the presence of solar thermal tides in VMC observations. Since only about 8 h of local time are covered from VMC measurements no reliable coefficients can be obtained from the employed model at this moment. We are continuing our efforts to expand the local time coverage of our tracking results to regions closer to the terminator regions.
5. Discussion and Conclusions [50] Figure 11 shows a comparison between zonal wind speed profiles from VMC with previous observations from Pioneer Venus and Galileo. Although maximum resolutions of Mariner 10 Images lie within the same range as VMC data, lack of sufficient overlap with required time interval prevented measurements from the high-resolution images. Therefore we did not include results from Mariner 10 wind tracking in this comparison. [51] The Pioneer Venus profile represents an average of the results from 1980 and 1982 imaging seasons of the OCPP polarimeter [Limaye, 2007] which is in good agreement with the VMC profile. The general shape of the Pioneer Venus orbit was quite similar to the one of Venus
Express, with a pericenter at about 15°N the planet was observed mainly from a near-equatorial perspective. In contrast to the Pioneer Venus orbit, the 80°N pericenter of Venus Express allows VMC to deliver improved coverage of the Southern Hemisphere in nadir geometry. The average resolution of the OCPP images used for cloud tracking also lies around 30 km/pixel. One significant limitation of the OCPP image data is the 4 h time interval between images [Rossow et al., 1980, 1990]. Whereas the instantaneous image acquisition of VMC allows for arbitrary intervals between images. [52] Also the results from the high-resolution cloud tracking in Galileo images are in good general agreement with our findings. Prior to the VMC observations the Galileo images comprised the best resolved (down to 15 km/pixel) sequences used for cloud tracking and have recently been reused for high-resolution wind tracking [Toigo et al., 1994; Peralta et al., 2007]. Time intervals between Galileo images range from 15 min to 2 h, thus having in principle the same temporal resolution as the VMC sequences. [53] So far, dynamical observations of the Venus clouds have delivered either only highly resolved short-term or broadly averaged long-term results. With VMC it finally became possible to investigate the global dynamics with a relatively high resolution in space and time on a long-term basis. Our findings from manual wind tracking in VMC UV image sequences so far show that the details of the cloud top level dynamics of Venus appear to be highly variable, down to a timescale of days. On the other side the average general rotation of the atmosphere remained effectively stable since Mariner 10, especially with regard to the remarkable agreement between VMC measurements and results from previous missions. [54] In some cases significant variations in the zonal wind properties were observed especially in the mid latitude regions on an orbit to orbit basis, indicating possible
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Figure 11. (a) Comparison of zonal wind speed profiles from VMC, Pioneer Venus, and Galileo measurements. (b) Corresponding rotation periods. influences from global-scale wave modes. In other cases we measured more constant conditions over one atmospheric revolution or even longer periods, at cloud top level. These short-term variations between results from individual orbit tracking sequences appear to have an impact on the average zonal wind component profiles over longer-term observation periods. So far the search for Kelvin gravity wave modes in the consecutive measurements from orbits 281 – 284 and 460– 463 did not yield any conclusive results. Most likely longer sequences of coherent tracking data will be required for clear identification of these global wave modes in VMC data. Also it remains an interesting question whether the irregularly detected mid latitude jets are indeed variable with time or simply become, in parts or in total, shielded from view by higher H2SO4 haze layers for varying time intervals. [55] Winds at latitudes higher than 60°S are still difficult to track because of low contrast and scarcity of features, but increasing data is being collected. Over all, it was possible to extend latitudinal coverage of the cloud top winds with VMC observations. [56] A first analysis indicated that diurnal and semidiurnal thermal tide modes appear to be present in the data, but a more profound confirmation still depends on synthesis of VIRTIS and VMC observations on night and dayside in order to bridge the large data gap in local time coverage of the VMC UV data between 17 and 9 h local solar time. [57] Although poorly resolved, results on average meridional wind speed components are in general agreement with previous observations and with the presence of a Hadley cell spanning between equatorial region and about 45°S latitude. Since measurement accuracy is very low so far, so we decided not to use the results on meridional winds for detailed investigations. We expect future data to improve the accuracy up to a degree that allows for more reliable meridional profiles fit for being used to draw more conclusions. [58] As VMC continues to acquire data, not only more UV data will become available for dayside cloud top level tracking purposes. In addition we are investigating whether data from the two IR channels of the instrument could be used for tracking lower-level features on the day side and
extend cloud tracking to the night side. Although contrasts range only from 1 – 5% for the IR features in day side images, it might be feasible to improve contrasts through carefully applied filtering processes. Tracking on the night side would prove even more difficult, since VMC is registering signals from the surface simultaneously with those from the lower cloud levels. Efforts for both tracking possibilities are in progress at the moment and will be presented in future works. [59] Meanwhile operations are increasingly adapted and custom fitted for wind tracking in the UV channel, improving steadily on sequence length and also on timing between individual images. Furthermore wind tracking sequences are extended down to lower altitudes above the planet to allow for higher resolutions and more accurate tracking. One method to achieve this goal is to follow cloud features to compensate for the small area covered by pericenter observations. We hope to achieve tracking sequences with resolutions as high as about 7 km/pixel. [60] Acknowledgments. We would like to thank A. Sanchez-Lavega and R. Hueso for their valuable cooperation and making their latest wind tracking results available to comparison for us. Thanks also go to the International Max Planck Research School on Physical Processes in the Solar System and beyond, the rest of the VMC team for making this work possible, and to NASA for contributions to this paper from grants NNG06GC68G and NNX07AF27G.
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Gierasch, P., et al. (1997), The general circulation of the Venus atmosphere: An assessment, in Venus II, edited by Stephen W. Bougher, D. M. Hunten, and R. J. Philips, pp. 459 – 500, Univ. of Ariz. Press, Tucson, Ariz. Kawabata, K., D. L. Coffeen, J. E. Hansen, W. A. Lane, M. Sato, and L. D. Travis (1980), Cloud and haze properties from Pioneer Venus polarimetry, J. Geophys. Res., 85, 8129 – 8140, doi:10.1029/JA085iA13p08129. Limaye, S. S. (1985), Venus atmospheric circulation: Observations and implications of the thermal structure, Adv. Space Res., 5(9), 51 – 62, doi:10.1016/0273-1177(85)90270-4. Limaye, S. S. (1988), Venus: Cloud level circulation during 1982 as determined from the Pioneer Cloud Photopolarimeter images. Part II. Solar longitude dependent circulation, Icarus, 73, 212 – 226, doi:10.1016/00191035(88)90094-2. Limaye, S. S. (2007), Venus atmospheric circulation: Known and unknown, J. Geophys. Res., 112, E04S09, doi:10.1029/2006JE002814. Limaye, S. S., and V. E. Suomi (1981), Cloud motions on Venus: Global structure and organization, J. Atmos. Sci., 38, 1220 – 1235, doi:10.1175/ 1520-0469(1981)038<1220:CMOVGS>2.0.CO;2. Limaye, S. S., C. J. Grund, and S. P. Burre (1982), Zonal mean circulation at the cloud level of Venus: Spring and fall 1979 OCPP observations, Icarus, 51, 416 – 439, doi:10.1016/0019-1035(82)90092-6. Limaye, S. S., C. Grassotti, and M. J. Kuetemeyer (1988), Venus: Cloud level circulation during 1982 as determined from the Pioneer Cloud Photopolarimeter images. Part I. Time and zonally averaged circulation, Icarus, 73, 193 – 211. Markiewicz, W. J., et al. (2007a), Venus Monitoring Camera for Venus Express, Planet. Space Sci., 55(12), 1701 – 1711, doi:10.1016/j.pss. 2007.01.004. Markiewicz, W. J., D. V. Titov, S. S. Limaye, H. U. Keller, N. Ignatiev, R. Jaumann, N. Thomas, H. Michalik, R. Moissl, and P. Russo (2007b), Morphology and dynamics of the upper cloud layer of Venus, Nature, 450, 633 – 636, doi:10.1038/nature06320. Peralta, J., R. Hueso, and A. Sanchez-Lavega (2007), A reanalysis of Venus winds at two cloud levels from Galileo SSI images, Icarus, 190, 469 – 477, doi:10.1016/j.icarus.2007.03.028. Piccialli, A., D. V. Titov, D. Grassi, I. A. Khatuntsev, P. Drossart, G. Piccioni, and A. Migliorini (2008), Cyclostrophic winds from the VIRTIS temperature sounding: A preliminary analysis, J. Geophys. Res., 113, E00B11, doi:10.1029/2008JE003127. Rossow, W. B., A. D. Del Genio, S. S. Limaye, L. D. Travis, and P. H. Stone (1980), Cloud morphology and motions from Pioneer-Venus images, J. Geophys. Res., 85, 8107 – 8128, doi:10.1029/ JA085iA13p08107.
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Rossow, W. B., A. D. Del Genio, and T. Eichler (1990), Cloud-tracked winds from Pioneer Venus OCPP images, J. Atmos. Sci., 47, 2053 – 2084, doi:10.1175/1520-0469(1990)047<2053:CTWFVO>2.0.CO;2. Sa´nchez-Lavega, A., et al. (2008), Variable winds on Venus mapped in three dimensions, Geophys. Res. Lett., 35, L13204, doi:10.1029/ 2008GL033817. Schubert, G. (1983), General circulation and the dynamical state of the Venus atmosphere, in Venus, edited by D. Hunten, L. Colin, T. Donahue, and V. Moroz, pp. 651 – 765, Univ. of Ariz. Press, Tucson, Ariz. Schubert, G., S. W. Bougher, C. C. Covey, A. D. Del Genio, A. S. Grossman, J. L. Hollingsworth, S. S. Limaye, and R. E. Young (2007), Venus atmosphere dynamics: A continuing enigma, in Exploring Venus as terrestrial planet, Geophys. Monogr. Ser., vol. 176, edited by L. W. Esposito, E. R. Stofan, and T. E. Cravens, pp. 121 – 138, AGU, Washington, D. C. Smith, M. D., and P. J. Gierasch (1996), Global-scale winds at the Venus cloud top inferred from cloud streak orientations, Icarus, 123(2), 313 – 323, doi:10.1006/icar.1996.0160. Suomi, V. E. (1975), Cloud motions on Venus, in The Atmosphere of Venus, edited by J. E. Hansen, 42 pp., NASA, Washington, D. C. Svedhem, H., et al. (2007), Venus Express—The first European mission to Venus, Planet. Space Sci., 55(12), 1636 – 1652, doi:10.1016/j.pss. 2007.01.013. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54(13 – 14), 1279 – 1297, doi:10.1016/j.pss.2006.04.017. Toigo, A., P. J. Gierasch, and M. D. Smith (1994), High resolution cloud feature tracking on Venus by Galileo, Icarus, 109(2), 318 – 336, doi:10.1006/icar.1994.1097. Tomasko, M. G., L. R. Doose, and P. H. Smith (1985), The absorption of solar energy and the heating rate in the atmosphere of Venus, Adv. Space Res., 5(9), 71 – 79, doi:10.1016/0273-1177(85)90272-8.
M. Almeida, European Space Astronomy Centre, European Space Agency, P.O. Box - Apartado de Correos 78, E-28691 Madrid, Spain. T. Behnke, R. Jaumann, K.-D. Matz, and T. Roatsch, Institut fur Planetenforschung, Deutsches Zentrum fur Luft- und Raumfahrt, Rutherfordstrasse 2, D-12489 Berlin, Germany. S. F. Hviid, W. J. Markiewicz, R. Moissl, G. Portyankina, and D. V. Titov, Max-Planck-Institut fuer Sonnensystemforschung, Max-PlanckStrasse 2, D-37191 Katlenburg-Lindau, Germany. (
[email protected]) N. I. Ignatiev and I. Khatuntsev, Space Research Institute, 84/32 Profsoyuznaya Street, Moscow 117997, Russia. S. S. Limaye, Space Science and Engineering Center, University of Wisconsin-Madison, 1017 Atmospheric Oceanic and Space Sciences Building, 1225 West Dayton Street, Madison, WI 53706, USA.
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Cyclostrophic winds from the Visible and Infrared Thermal Imaging Spectrometer temperature sounding: A preliminary analysis A. Piccialli,1 D. V. Titov,1,2 D. Grassi,3 I. Khatuntsev,2 P. Drossart,4 G. Piccioni,5 and A. Migliorini5 Received 25 February 2008; accepted 25 July 2008; published 22 October 2008.
[1] Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) onboard the Venus
Express spacecraft has been operating since April 2006 providing new observations of the temperature structure of Venus mesosphere (60–95 km). Zonal winds in the middle atmosphere of Venus have been retrieved from VIRTIS temperature profiles using the cyclostrophic approximation. The wind field is characterized by three main features: (1) a midlatitude jet connected to the thermal feature known as the cold collar; the jet reaches a maximum speed of 80–90 ± 10 m/s near the cloud top (70 km altitude) at 50°S latitude; (2) a strong decrease of wind from 60°S toward the pole reaching zero velocity at 70°S; and (3) the decrease of the wind above the jet with increasing altitude. Local time dependence of the temperature and wind field has been analyzed. Temperatures show at cloud tops a cooling of 15 K during the night which propagates also on the zonal wind field. Comparison with cloud-tracked winds from Venus Monitoring Camera images indicates a first-order agreement. Citation: Piccialli, A., D. V. Titov, D. Grassi, I. Khatuntsev, P. Drossart, G. Piccioni, and A. Migliorini (2008), Cyclostrophic winds from the Visible and Infrared Thermal Imaging Spectrometer temperature sounding: A preliminary analysis, J. Geophys. Res., 113, E00B11, doi:10.1029/2008JE003127.
1. Introduction [2] Early observations showed that the Venus atmosphere has two different dynamic regimes. In the thermosphere above 100 km diurnal temperature gradients force global solar-antisolar circulation [Bougher et al., 1997; Lellouch et al., 1997]. Troposphere and mesosphere (0 – 100 km) are involved in the retrograde almost purely zonal motion with maximum wind speed of about 100 m/s at the cloud tops, called superrotation. The physical mechanisms that maintain this global regular motion are still not known [Schubert, 1983; Gierasch et al., 1997; Schubert et al., 2007]. Leovy [1973] was first to notice that on a slowly rotating planet, where Coriolis force is negligible, the global zonal circulation is a result of local balance of pressure gradient and centripetal force which is called cyclostrophic balance. The thermal wind equation that describes this balance relates vertical shear of zonal wind to latitudinal temperature gradient on a constant pressure surface [Leovy, 1973; Schubert, 1983]. Although the cyclostrophic balance seems to clearly describe the observed state of Venus superrotation, it does not explain what originally brought the atmo-
1 Max Planck Institute for Solar System Research, Katlenburg-Lindau, Germany. 2 Space Research Institute, Moscow, Russia. 3 Istituto di Fisica dello Spazio Interplanetario, INAF, Rome, Italy. 4 LESIA, Observatoire de Paris, Meudon, France. 5 Istituto di Astrofisica Spaziale e Fisica Cosmica, INAF, Rome, Italy.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003127$09.00
sphere to this state or which mechanisms maintain the vertical wind shear. [3] Temperature structure of the Venus mesosphere was sounded remotely by radio occultation and thermal emission radiometry and spectroscopy on the Pioneer Venus and Venera-15 orbiters [Seiff, 1983; Taylor et al., 1983; Seiff et al., 1985; Lellouch et al., 1997; Zasova et al., 2007]. The mesospheric temperature field showed two remarkable peculiarities. Temperatures above 75 km monotonically increase from equator to pole thus creating positive latitudinal temperature gradient at constant pressure level. Inside the upper cloud layer (60 –75 km) the trend reverses and temperature field forms a bulge of cold air at 60° – 80° latitude which was called ‘‘cold collar.’’ [4] The measured temperatures were used to derive zonal wind field using the cyclostrophic approximation [Chub and Yakovlev, 1980; Seiff, 1983; Newman et al., 1984; Limaye, 1985; Roos-Serote et al., 1995; Zasova et al., 2007]. These studies constrained the winds in the poorly explored region of transition to the solar-antisolar thermospheric circulation. The main features of the cyclostrophic wind field are (1) the presence of strong midlatitude jet centered at 50° latitude and 70 km altitude and (2) the decrease of zonal wind with altitude above the cloud top. This behavior is governed by the latitudinal temperature contrasts. In the midlatitudes the atmosphere is accelerated at the cloud level by negative latitudinal gradient of temperature associated with the cold collar. Above the cloud tops on the contrary positive temperature gradient forces the winds to fade out with altitude. The cyclostrophic balance was found to break above 70– 75 km in high (>70°) and low
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Figure 1. Contours of temperature field averaged over the range 2000 – 2200 LT for the VIRTIS-M orbit VI0072. Contour interval is 5 K. In gray the cloud layer is represented schematically. (<40°) latitudes. Here the latitudinal temperature gradient cannot provide centripetal force for the air parcels involved in the zonal motion. The magnitude of the midlatitude jet changes from 160 m/s in the work by Newman et al. [1984] to 100 m/s in the work by Zasova et al. [2007]. Position of the jet core showed meandering in the latitude range of 45° – 65° and altitudes of 65– 70 km. Zasova et al. [2007] also reported about diurnal and semidiurnal harmonics present in the thermal wind pattern derived from Venera15 data. [5] Since April 2006, Venus Express carries out systematic remote sensing observations of the Venus mesosphere using both radio occultation and thermal emission spectroscopy techniques [Svedhem et al., 2007; Ha¨usler et al., 2006; Drossart et al., 2007]. Polar orbit of the satellite allows the experiments to achieve full latitude coverage and especially focus of the southern hemisphere poorly studied in the earlier missions [Titov et al., 2006]. Here we present the first results of calculations of the cyclostrophic wind field from the temperature sounding by VIRTIS experiment [Grassi et al., 2008].
2. Temperature Field in the Mesosphere [6] Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) is one of the experiments on board of Venus Express. It consists of two channels. VIRTIS-M is an imaging spectrometer with moderate spectral resolution (R200) and high spatial resolution (0.25 mrad) operating in the spectral ranges 0.25 – 1 mm and 1 –5 mm. VIRTIS-H is a high-resolution spectrometer (R1200) operating in the spectral range 1.84– 4.99 mm [Drossart et al., 2007; Piccioni et al., 2008]. VIRTIS-M acquires thermal emission spectra in the region of 4.3 mm CO2 band. These measurements were then used to retrieve vertical profiles of air temperature at 67 pressure levels in the Venusian mesosphere. The VIRTIS temperature sounding covers altitude range from 85 to 65 km (1 – 100 mbar). The overall error in retrieved temperature is<4 K in the range 100 –0.1 mbar and
is better than 1 K between 70 and 7 mbar. The main sources of errors and systematic uncertainties in the temperature retrievals are instrumental noise in the spectral ranges sounding high altitudes and uncertainties in the aerosols densities within the cloud deck (70 – 60 km) [Grassi et al., 2008]. The VIRTIS observations analyzed in this work cover the night side in the southern hemisphere. The measurements on the day side have nonnegligible solar component and were excluded from our study. [7] Figures 1 and 2 show example of the latitude-altitude temperature field retrieved from the VIRTIS-M observations in orbit 72 and corresponding latitude dependence of temperature at constant pressure levels. We used Chebyshev polynomials of degree 4 to fit the data points and to evaluate the latitudinal temperature gradient. Above 75 km (20 mbar) temperatures on isobaric surfaces generally increase toward the pole. Below this level the temperature pattern becomes more complex and at 50°– 70° latitude shows the region of temperature inversions at the cloud top (cold collar). In Figure 2 the cold collar is clearly visible as dip between 50° and 70° latitude at 40 and 90 mbar. Both positive temperature gradient in the higher mesosphere and cold collar right at the cloud tops were observed by Pioneer Venus [Taylor et al., 1983; Newman et al., 1984] and Venera-15 [Zasova et al., 2007] missions in the northern hemisphere. VIRTIS observations in the south strongly suggest global hemispheric symmetry of the temperature field.
3. Zonal Thermal Winds [8] Following Newman et al. [1984] we used cyclostrophic approximation to derive zonal wind pattern in the mesosphere from the VIRTIS temperature field. If we indicate with u the zonal wind speed in m/s, with z = ln (p/pref) the logarithmic pressure vertical coordinate, with R the gas constant and with (@T/@f)p = const the latitudinal temperature gradient, it is possible to describe the zonal wind u by the thermal wind equation [Schubert, 1983]
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Figure 2. Plots of temperature at different pressure levels for 2000 –2200 LT. Chebyshev polynomial approximation is represented in gray. Error bar shows random error of air temperature retrieval. Data refer to the VIRTIS-M orbit VI0072. @u R @T 2u ¼ @z tan f @f p¼const
ð1Þ
The retrieved air temperatures T(f) were approximated by Chebyshev polynomials of degree 4 between 25° and 75° of latitude at 49 pressure levels. The lower boundary condition needed to solve differential equation (1) was
Figure 3. Functions used as lower boundary condition. Curve 1 is the function described by equation (2) used for the nominal case. Curve 2 is the solid body rotation function u0 = 90 cos f; as can be seen, it perfectly fits the curve 1 for latitudes higher than 70°, while at the equator it results 10 m/s stronger. Curve 3 represents the cloud-tracked wind derived from Galileo SSI NIR images [Peralta et al., 2007]. This curve has the same magnitude of curve 1 near the equator, but a strong decrease in the wind speed is seen at latitudes higher than 40°. 3 of 9
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Figure 4. Contours of the difference absolute value (m/s) between the zonal thermal wind speed derived from VIRTIS temperature retrievals assuming u0 = (45 sec h((f 56)/9) + 75) cos f as lower boundary condition and the zonal thermal wind derived assuming (a) u0 = 90 cosf and (b) cloud-tracked winds derived from Galileo SSI NIR images as lower boundary condition. Data refer to the VIRTIS-M orbit VI0072. taken at the reference pressure level pref = 275 mb (58 km) using the equation adopted by Counselman et al. [1980]; coefficients were selected to fit the Venus
Monitoring Camera (VMC) direct measurements of wind profile and were provided by I. Khatuntsev (personal communications, 2008):
Figure 5. Latitude dependence of temperatures at a pressure level of 90 mb for 2000 – 2200 LT. Chebyshev polynomial approximation is represented in black together with the ±1s, ±2s curves represented in gray. Error bar shows random error of air temperature retrieval. Data refer to the VIRTISM orbit VI0072. 4 of 9
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Figure 6. Contours of the zonal thermal wind speed (m/s) for 2000– 2200 LT derived from VIRTIS temperature retrievals assuming cyclostrophic balance. Contour interval is 5 m/s. Data refer to the VIRTIS-M orbit VI0072. In gray the cloud layer is represented schematically.
U0 ðfÞ ¼ ½45 sech ððf 56Þ=9Þ þ 75 cos f
ð2Þ
In order to test our retrieval code we used temperature field retrieved from Venera-15 Fourier Spectrometer (FS)
data [Zasova and Khatuntsev, 1997]. Our test calculations were in good agreement with original results. [9] We studied sensitivity of calculated thermal winds to numerical parameters and lower boundary conditions. First, we found that the Chebyshev polynomial of degree 4
Figure 7. Plots of latitude dependence of temperatures at selected pressure levels; different colors correspond to different local time: (blue) 1800– 2000 LT; (green) 2000– 2200 LT; (yellow) 2200 –2400 LT; (red) 0000 –0200 LT; (light blue) 0200 – 0400 LT; and (orange) 0400– 0600 LT. Data refer to the VIRTIS-M orbit VI0072. 5 of 9
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Figure 8. Contours of the zonal thermal wind speed (m/s) for (a) 1800 – 2000 LT and (b) 0300 –0500 LT. Winds have been derived from VIRTIS temperature retrievals assuming cyclostrophic balance. Contour interval is 5 m/s. Data refer to 13 VIRTIS orbits acquired between May and December 2006. provided the best fit to the data (Figure 2). Higher power of polynomial did not improve the data fit and resulted in highfrequency oscillations. Second, we considered different functions as lower boundary condition (Figure 3). [10] We calculated the difference between the zonal wind field derived from VIRTIS temperature profiles using curve 1 as lower boundary condition and that obtained using curves 2 and 3 (Figures 4a and 4b). Both cases show small discrepancy with the nominal case (equation (2)) confirming that the retrieved wind weakly depends on the lower boundary condition. In Figure 4a the difference reaches maximum value of 40 m/s because of the absence in curve 2 of the midlatitude bulge present in curve 1 that forces winds to slow down with altitude faster than in the nominal case. In Figure 4b a bigger discrepancy is observed caused by the strong wind decrease at midlatitudes seen in Galileo NIR images. Both test cases show that lower boundary condition does not affect the region of the jet. [11] Errors in the temperature retrievals are the source of uncertainty on the derived wind speed. We followed the approach used by Newman et al. [1984] to assess the propagation of temperature retrieval error on the wind field.
On every pressure level air temperatures were averaged within 5° latitude bins. Mean value and standard deviation of each data point from approximating Chebyshev polynomial were calculated and centered in the interval. Another Chebyshev polynomial was used to fit the standard deviations and was added to the original polynomial approximation curve to produce the +1s curve. The 1s and ±2s curves were calculated in a similar way. Figure 5 shows that almost all the data points are within ±2s range. Since scattering of the retrieved temperatures depends on latitude the gradients of ±1s, ± 2s curves differ from that of the original approximation curve that results in distortions of the wind field. [12] Wind speeds were retrieved from ±1s, ±2s curves, using equation (1) and the boundary conditions from equation (2) to assess the error in wind velocity determination. For all the curves, the position of the jet is not changed within 2° latitude. The wind field calculated for 1s curve has a midlatitude jet speed of 89 m/s at 51° latitude. The speed of the jet for the 2s curve is slightly altered. On the other hand for the 1s and 2s curves the speed of the jet reaches a value of 98 and 102 m/s
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Figure 9. (a) Latitude of midlatitude jet versus local time for thirteen VIRTIS orbits acquired between May and December 2006; (b) Altitude of midlatitude jet versus local time; (c) Wind speed of midlatitude jet versus local time.
respectively. As an esteem of retrieval error on the midlatitude jet speed, it seems reasonable to take a value of ± 10 m/s which is comparable to the uncertainties on cloud-tracked winds derived from the Venus Monitoring Camera (R. Moissl et al., Cloud top winds from tracking UV features in VMC images, submitted to Journal of Geophysical Research, 2008). [13] Figure 6 shows an example of thermal wind for 2000 – 2200 local time (LT) derived from the VIRTIS-M observations in orbit 72. The main feature in the plot is the midlatitude jet centered at 50° latitude and 67 km of altitude with a maximum speed of 90 m/s. Comparison of the wind field (Figure 6) with the temperature field (Figure 1) shows that the midlatitude jet is related to the cold collar. [14] The thermal feature known as the warm polar mesosphere, characterized by an increase of temperature toward the pole on isobaric surfaces between 75 and 90 km of altitude, forces the retrieved zonal wind to decrease to zero very fast at higher latitudes. It is important to note, however, that for latitudes lower than 30° and higher than 75° the cyclostrophic balance fails, thus wind retrievals at these latitudes should not be taken in account.
[15] The VIRTIS observations in orbit 72 completely covered the night side of the southern hemisphere that allowed us to estimate the dependence of temperature field on local time. Figure 7 shows latitude profiles of temperatures at selected pressure levels for 2-h local time bins. They suggest that the atmosphere at the cloud tops cools down by about 15 K during the night. Theses changes in temperature structure also affect the thermal wind field. [16] An example of the effect of diurnal variations of temperature structure is shown in Figures 8a and 8b. Wind field has been retrieved for two different local times combining the temperature profiles of thirteen VIRTIS orbits acquired between May and December 2006. [17] Last, we examined in detail the dependence of the midlatitude jet on local time. Temperature profiles have been divided in local time intervals of 3 h and the wind speed has been retrieved for each local time bin. The behavior of midlatitude jet position and speed with local time has been analyzed in Figures 9a – 9c. A change in the position and speed of the midlatitude jet can be clearly observed, however the magnitude of the variations lays within the uncertainties on retrieved wind speeds. There-
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fore, any conclusion on local time dependence should be taken cautiously.
4. Discussion and Conclusions [18] We presented the preliminary results of the cyclostrophic wind calculations from the VIRTIS-M temperature sounding in the southern hemisphere on the night side [Grassi et al., 2008]. The main features of the wind field are (1) the presence of the midlatitude jet with peak velocity of 80 –90 m/s centered at 50°S at the cloud tops (70 km); (2) fast decrease of zonal winds poleward from 60°S with zero velocity reached at 70°S; and (3) gradual decrease of thermal wind with altitude above the jet. Our results are in general agreement with those based on the earlier observations [Newman et al., 1984; Zasova et al., 2007]. These features correlate with the behavior of the temperature field (Figure 1). According to the thermal wind equation (1) the negative latitude gradient within the upper cloud (<70 km) associated with the cold collar accelerates the wind. This trend changes with reversal of the temperature gradient above the cloud that causes thermal winds to fade out toward the top of the mesosphere. This deceleration is the strongest in high latitudes (>65°S) that results in that cyclostrophic balance breaks down in polar mesosphere (Figure 6). [19] The magnitude of the jet in our results (80 – 90 m/s) is smaller than that reported by Zasova et al. [2007] (90 – 110 m/s) and especially by Newman et al. [1984] (140 – 160 m/s). This discrepancy can be partially attributed to the differences in numerical schemes and boundary conditions used in calculations, but seems to mainly result from peculiarities of temperature sounding techniques. Radio occultation data used by Newman et al. [1984] provide temperature structure with vertical resolution of few hundred meters that allows the measurements to completely resolve deep temperature inversions typical for the cold collar regions. Vertical resolution in thermal emission spectroscopy in the CO2 bands used by Zasova et al. [2007] and in this work does not exceed few kilometers. This smoothes temperature inversions and effectively reduces the latitudinal gradient of temperature that eventually accelerates the wind. [20] Comparison of the thermal wind field to the cloudtracked winds is the test bench for cyclostrophic balance assumption. Results of our calculations in first order agree with the Venus Express measurements of the cloud top winds [Sanchez-Lavega et al., 2008; R. Moissl et al., submitted manuscript, 2008] that gave the wind speed of 80– 100 m/s at 50°S and fast decrease of zonal wind poleward. However, our thermal wind deviates from the observed one in low latitudes (<40°S), indicating that cyclostrophic balance is not valid here. The core of the cyclostrophic jet is located almost exactly at the cloud tops (70 km) in all thermal wind calculations. We note, however, that the latitudinal profiles of thermal wind at this altitude have much sharper maxima than those observed in cloud-tracked winds. [21] Following the earlier studies, especially the one by Zasova et al. [2007], we searched for local time variability in the properties of the midlatitude jet. The temperature field clearly indicates radiative cooling by 15 K of the night side atmosphere at the cloud tops (Figure 7) which also
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propagates to the thermal wind field (Figures 9a – 9c). However, we prefer to be cautious in making conclusions on this basis. First, the coverage by VIRTIS temperature sounding that we used so far is limited by the night side. And, second and probably more important, the weak meandering of the thermal wind field of 10– 20 m/s that results from diurnal variations of temperature structure seems to be within the ‘‘ignorance’’ range of the cyclostrophic hypothesis itself. For instance, the equation (1) was derived ignoring the meridional wind component which is of 15 m/s. So there could be some doubt that the wind field variations of 10 – 20 m/s, although derived by correct numerical procedures, are physically meaningful. [22] The future work will include the use of VIRTIS temperature sounding extended to the northern hemisphere and the day side. This would allow us to better constrain the thermal wind at 20°– 40°S and to extend the coverage to the northern hemisphere and to study hemispheric symmetry of the wind field. Days side temperature sounding would be needed to complete local time coverage and would allow comparison with simultaneously derived cloud-tracked winds. And, finally, extensive radio occultation sounding by VeRa/Venus Express would allow the thermal wind calculations to be extended to as deep as 40 km. Comparison of the thermal wind field derived from the combined VIRTIS/VeRa temperature structure with the winds derived from cloud features tracking in UV (70 km) and near IR on the night side (50 km) would eventually allow one to conclude about the regions of validity of the cyclostrophic balance. [23] Acknowledgments. We thank S. Limaye for fruitful discussions of the results. A. Piccialli is grateful to the Max Planck Research School for providing the opportunity to carry out this study. VIRTIS-Venus Express is an experiment developed jointly by IASF-INAF (Italy) and LESIA, Observatoire de Paris (France). The project is funded by ASI and CNES.
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Limaye, S. S. (1985), Venus atmospheric circulation: Observations and implication of the thermal structure, Adv. Space Res., 5, 51 – 62, doi:10.1016/0273-1177(85)90270-4. Newman, M., G. Schubert, A. J. Kliore, and I. R. Patel (1984), Zonal winds in the middle atmosphere of Venus from Pioneer Venus radio occultation d a t a , J . A t m o s . S c i . , 4 1 , 1 9 0 1 – 1 9 1 3 , d o i : 1 0 . 11 7 5 / 1 5 2 0 0469(1984)041<1901:ZWITMA>2.0.CO;2. Peralta, J., R. Hueso, and A. Sa´nchez-Lavega (2007), A reanalysis of Venus winds at two cloud levels from Galileo SSI images, Icarus, 190(2), 469 – 477, doi:10.1016/j.icarus.2007.03.028. Piccioni, G., et al. (2008), VIRTIS (Visible and Infrared Thermal Imaging Spectrometer) for Venus Express, Eur. Space Agency, Noordwijk, Netherlands, in press. Roos-Serote, M., P. Drossart, T. Encrenaz, E. Lellouch, R. W. Carlson, K. H. Baines, F. W. Taylor, and S. B. Calcutt (1995), The thermal structure and dynamics of the atmosphere of Venus between 70 and 90 km from the Galileo-NIMS spectra, Icarus, 114(2), 300 – 309, doi:10.1006/ icar.1995.1063. Sanchez-Lavega, A., et al. (2008), Variable winds on Venus mapped in three dimensions, Geophys. Res. Lett., 35, L13204, doi:10.1029/ 2008GL033817. Schubert, G. (1983), General circulation and the dynamical state of the Venus atmosphere, in Venus, edited by D. M. Hunten, et al., pp. 651 – 765, Univ. of Ariz. Press, Tucson. Schubert, G., S. W. Bougher, C. C. Covey, A. D. Del Genio, A. S. Grossman, J. L. Hollingsworth, S. S. Limaye, and R. E. Young (2007), Venus atmosphere dynamics: A continuing enigma, in Exploring Venus as a Terrestrial Planet, Geophys. Monogr. Ser., vol. 176, edited by L. W. Esposito, E. R. Stofan, and T. E. Cravens, pp. 121 – 138, AGU, Washington, D. C. Seiff, A. (1983), Thermal structure of the atmosphere of Venus, in Venus, edited by D. M. Hunten, et al., pp. 215 – 279, Univ. of Ariz. Press, Tucson.
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Seiff, A., J. T. Schofield, A. J. Kliore, F. W. Taylor, S. S. Limaye, H. E. Revercomb, L. A. Sromovsky, V. V. Kerzhanovich, V. I. Moroz, and M. Y. Marov (1985), Models of the structure of the atmosphere of Venus from the surface to 100 kilometers, Adv. Space Res., 5, 3 – 58, doi:10.1016/0273-1177(85)90197-8. Svedhem, H., et al. (2007), Venus Express—The first European mission to Ve n u s , P l a n e t . S p a c e S c i . , 5 5 , 1 6 3 6 – 1 6 5 2 , d o i : 1 0 . 1 0 1 6 / j.pss.2007.01.013. Taylor, F. W., D. M. Hunten, and L. V. Ksanfomaliti (1983), The thermal balance of the middle and upper atmosphere of Venus, in Venus, edited by D. M. Hunten, et al., pp. 650 – 680, Univ. of Ariz. Press, Tucson. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017. Zasova, L. V., and I. V. Khatuntsev (1997), Thermal zonal wind in the Venus middle atmosphere according to Venera 15 IR-spectrometry, Adv. Space Res., 19, 1181 – 1190, doi:10.1016/S0273-1177(97)00269-X. Zasova, L. V., N. Ignatiev, I. Khatuntsev, and V. Linkin (2007), Structure of the Venus atmosphere, Planet. Space Sci., 55, 1712 – 1728, doi:10.1016/ j.pss.2007.01.011.
P. Drossart, LESIA, Observatoire de Paris, Section de Meudon, 5 place Jules Janssen, F-92195 Meudon, France. D. Grassi, Istituto di Fisica dello Spazio Interplanetario, INAF, Via del Fosso del Cavaliere 100, I-00133 Rome, Italy. I. Khatuntsev, Space Research Institute, 84/32 Profsojuznaja Str, 117997 Moscow, Russia. A. Migliorini and G. Piccioni, Istituto di Astrofisica Spaziale e Fisica Cosmica, INAF, Via del Fosso del Cavaliere 100, I-00133 Rome, Italy. A. Piccialli and D. V. Titov, Max Planck Institute for Solar System Research, Max Planck Strasse 2, D-37191 Katlenburg-Lindau, Germany.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114, E00B32, doi:10.1029/2008JE003073, 2009
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Gravity waves in the upper atmosphere of Venus revealed by CO2 nonlocal thermodynamic equilibrium emissions R. F. Garcia,1 P. Drossart,2 G. Piccioni,3 M. Lo´pez-Valverde,4 and G. Occhipinti5 Received 10 January 2008; revised 29 December 2008; accepted 14 January 2009; published 17 March 2009.
[1] The imaging capabilities of the Visible and Infrared Thermal Imaging Spectrometer-
Mapper (VIRTIS-M) onboard Venus Express mission are used to analyze perturbations of CO2 nonlocal thermodynamic equilibrium emissions in the thermosphere of Venus. These emissions with a wavelength of 4.3 mm originate from the 110–140 km altitude range and are sensitive to density perturbations. They show wave-like perturbations of about 0.5% root-mean-square amplitude of the background signal with horizontal wavelengths in the 90–400 km range. The horizontal phase velocities are similar in magnitude and direction from one orbit to the next, with averages of 70 m/s westward and 30 m/s northward. The geographical wave distribution and the orientation of wavefronts demonstrate that the polar vortex at work in the cloud layer is the source of these gravity waves. The large westward zonal phase velocity in the 0900–1500 local time range argues in favor of the superrotation dynamics within 30° from the south pole to an altitude of at least 115 km. The gravity wave dispersion relation and the geographical distribution of wavefront amplitudes suggest the presence of a solar to antisolar wind close to the south pole. Because the centroid altitude and the sensitivity of these emissions to gravity wave perturbations are only roughly estimated, it is not possible to make a quantitative estimate of the upward energy transfer. However, this study demonstrates the strong influence of the polar vortex on the circulation in the atmosphere of Venus up to the thermosphere. Citation: Garcia, R. F., P. Drossart, G. Piccioni, M. Lo´pez-Valverde, and G. Occhipinti (2009), Gravity waves in the upper atmosphere of Venus revealed by CO2 nonlocal thermodynamic equilibrium emissions, J. Geophys. Res., 114, E00B32, doi:10.1029/2008JE003073.
1. Introduction [2] The atmosphere of Venus is a very active and turbulent medium which is expected to show competition between subsolar to antisolar and superrotation circulations [Gierasch et al., 1997]. The clues to understanding the mechanisms of this competition are in the lower thermosphere of Venus. However, up to now, the dynamics of Venus’ lower thermosphere have only been investigated by indirect evidence relying mainly on the distribution of neutral species, temperature and O2 airglow [Bougher and Borucki, 1994; Bougher et al., 1997]. The nonlocal thermodynamic equilibrium (non-LTE) emissions of CO2 on the dayside upper atmosphere of Venus provide a way of investigating the structure and dynamics of Venus’ 1 Laboratoire de Dynamique Terrestre et Plane´taire, UMR5562, Observatoire Midi-Pyre´ne´es, Universite´ de Toulouse, CNRS, Toulouse, France. 2 LESIA, Observatoire de Paris, UPMC, Universite´ Paris-Diderot, CNRS, Meudon, France. 3 INAF, IASF, Rome, Italy. 4 Instituto de Astrofı´sica de Andalucı´a, CISC, Granada, Spain. 5 Institut de Physique du Globe de Paris, Universite´ Denis Diderot Paris 7, Paris, France.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003073$09.00
lower thermosphere [Rolda´n et al., 2000; Gilli et al., 2009]. Moreover, the imaging capabilities of the Visible and Infrared Thermal Imaging Spectrometer-Mapper (VIRTIS-M) onboard the Venus Express mission can be used to investigate the evolution over time of these emissions on successive images taken at nadir or at the limb [Drossart et al., 2004, 2007]. [3] The models of thermosphere circulation [Zhang et al., 1996; Bougher et al., 1997] are sensitive to the dynamics and breaking of atmospheric gravity waves. These waves have been observed in Venus troposphere [Gierasch, 1987; Young et al., 1987; Hinson and Jenkins, 1995] and thermosphere [Kasprzak et al., 1993] by different exploration missions. The modeling of the upper atmosphere gravity wave propagation suggests a source in the lower atmosphere [Mayr et al., 1988] which is probably related to the turbulence in the cloud layer [Woo et al., 1980, 1982]. Gravity waves are also used to explain the atmospheric superrotation [Alexander, 1992]. [4] The CO2 non-LTE emissions, which are mainly sensitive to density perturbations [Rolda´n et al., 2000], provide a unique opportunity for direct observation of the dynamics of gravity waves in the 110– 140 km altitude range. Despite the limitations induced by the integrated information content along the line of sight, and the restricted knowledge on both the space/time distribution of these emissions [Drossart et al.,
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Figure 1. (left) Limb emission profiles at 4.277 mm (solid line) and 4.315 mm (dashed line) on orbit 315. (right) Spectral profiles at 132 km and 118 km altitudes on the limb profile (vertical bar indicates the position of the wavelength at 4.277 mm). The spectral profiles are shifted by about 0.007 mm relative to values indicated in the text owing to a recent correction of spectral registration implemented in the text but not shown. 2007] and their sensitivity to density perturbations, the perturbations of these emissions can be used to characterize the wave content and dynamics of the upper atmosphere and its background circulation. [5] This study first describes the structure of CO2 nonLTE emissions at the 4.3 mm wavelength as seen by the VIRTIS-M spectral imaging instrument. The structure and dynamics of wave-like perturbations are then extracted through a specific data processing scheme. Finally, the source of wave excitation and the consequences on Venus circulation models are discussed.
2. CO2 Non-LTE Emissions at 4.3 mm [6] The non-LTE CO2 emissions in the upper atmosphere of Venus are described in detail by Gilli et al. [2009] in this issue. These emissions are due to the excitation of CO2 molecules by dayside solar radiations. Owing to the low atmospheric density above 90 km altitude, these molecules return to their ground state in a noncollisional manner by releasing photons in a series of emission bands around the 4.3 mm wavelength. Non-LTE situations of pure radiative excitation, like the present CO2 emissions, are relatively insensitive to the actual temperature, while their density dependence is more significant [Rolda´n et al., 2000; LopezPuertas et al., 2000]. In this section, we will recall only the main features of these emissions as seen by VIRTIS-M. [7] At the spectral resolution of VIRTIS-M, the limb profiles present two main peaks at 4.277 mm and 4.315 mm in the altitude range 110 to 140 km. The first emission peak is weaker than the second and presents a maximum on limb data at higher altitudes. The maximum emission strength and the altitude of the emission centroid vary with solar zenith angle and local time. The centroids of the emissions seen at the limb are in the 125–135 km range at the 4.277 mm wavelength and the 115–120 km range at 4.315 mm. An example of these emissions measured by VIRTIS-M at the
limb is shown in Figure 1 for the two emission wavelengths at 1300 local time (LT) and at 50.8° solar zenith angle on orbit 315. In Figure 1, the spectra around the 4.3 mm wavelength clearly show the two emission peaks at 4.277 mm and 4.315 mm plus additional non-LTE emissions above 4.4 mm. [8] When observed at nadir, CO2 non-LTE emissions at 4.3 mm show a clear variation with solar zenith angle, as can be seen in Figure 2 from a south polar view acquired during orbits 473 and 474. Some wavelike features, hardly discernible in Figures 2a and 2b, appear over the background emission. The rest of this study will discuss the structure and dynamics of these features. The structure of the background emission is not investigated.
3. Characteristics of Wave Patterns 3.1. Data Processing [9] This study uses the calibrated VIRTIS-M data for which continuum effects have been removed and absolute spectral radiances are expressed in W/(m2 mm sr). These data are extracted from the VIRTIS database and processed in the following steps: (1) correction of residual instrument discrepancies between odd and even samples of VIRTIS-M measurement line by an algorithm computing a smooth correction factor for each sample, (2) selection of image parts far from the limb to avoid limb effects, (3) enhancement of wave-like features through a filtering process based on the removal of low-pass filtered background emission, (4) projection of the filtered image into a geographical coordinate system centered on the subspacecraft point to avoid image distortions, (5) computation of wave displacements between two consecutive images for the same wavelength through an algorithm based on cross correlation, and (6) back-projection of the images and estimated displacement and velocity vectors into the true geographical system. [10] Calibrated VIRTIS-M data are corrected for instrument artifacts such as discrepancies between odd and even
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Figure 2. (a and b) Venus south polar views of radiance at the 4.315 mm wavelength averaged over four VIRTIS-M bands, for images taken along orbits 473 (Figure 2a) and 474 (Figure 2b). Color scale from black to white ranges from 0 to 0.06 W/(m2m msr). (c and d) Same images after wave enhancement by the data processing scheme. Color scales range from 1.5 to +1.5% of the background signal on the left (Figure 2c) and from 3 to +3% on the right (Figure 2d). Latitude contours are indicated in degrees, and the terminator line is almost horizontal. samples along the VIRTIS measurement line. However, owing to the low amplitudes of the wave signals investigated here, an additional correction of this instrument effect is applied in the first step of the algorithm. It is based on the idea that each pixel value at one wavelength pij is related to the true pixel radiance Rij by a correction factor cij in the following way: pij = Rijcij. This correction factor is estimated under the following hypotheses. cij is smooth along the sample direction (j) because the correction of a given sample is approximately the same for two consecutive measurement lines; and cij is close to 1 because the relative difference between odd and even samples of VIRTIS-M images is below 10% for more than 95% of the pixels. An algorithm is implemented on the basis of these two assumptions and the p true radiance is computed by Rij = cijij . [11] The second step is carried out manually to exclude regions of the image outside the 50° arc limit around the subspacecraft point. The third step is designed to enhance the wave features on nadir images at the 4.3 mm wavelength. In order to improve the signal-to-noise ratio, the images are first averaged over four VIRTIS-M wavelengths around the first and the second emission peaks (respectively from 4.26 mm to 4.29 mm and from 4.305 mm to 4.335 mm). The background radiation Bij is estimated by low-pass filtering of the radiance values Rij with a median filter 16 pixels wide. The filtered image is then computed through Fij = Rij Bij, and the normalized filtered image is estimated through Nij = Fij/Bij. The wave features in these images taken far from the planet, at spacecraft altitudes greater than 20,000 km, have
been shown to be almost insensitive to the median filter window for widths larger than 16 pixels. However, even if the use of a median filter is justified to reduce data spike effects, its frequency response is difficult to estimate. The normalized filtered images Nij are used to estimate the amplitudes of the waves relative to the background. Two examples of wave enhancement are presented for CO2 non-LTE emissions at 4.315 mm in Figures 2c and 2d during orbits 473 and 474. On the nightside, the low level of background signal creates spurious features not shown here. On the dayside, wavefront like features are visible which correspond to the small-scale features observed on the original images. [12] The projection and back-projection at steps 4 and 6 are performed to avoid distortions of pixels when plotted in geographical coordinates close to the south pole. The average subspacecraft point during the image acquisition is computed, and the pixels coordinates are simply rotated and back-rotated from the true geographical coordinates to a geographical coordinate system in which this point has zero latitude and longitude. In this manner, the latitudes and longitudes of pixels are kept within a range of ±30°. The images are projected on a regular (latitude, longitude) grid of 0.15° resolution in this new coordinate system. [13] When the VIRTIS data include two images of the same region at different times along the same orbit, the horizontal displacement of the waves can be estimated by cross correlation of the two images. The typical time difference between two consecutive images is about 1800 s. Similarly, if two wavelengths sample different altitudes, the
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Figure 3. Focus on two consecutive VIRTIS-M images at 4.315 mm wavelength filtered and projected at nadir along orbits (a and c) 473 and (b and d) 474. Color scales range from 1.5 to +1.5% of the background signal on the left and from 3 to +3% on the right. Black grid lines are plotted with a step of 25 pixels equal to the size of the correlation window. apparent displacement of the wavefront between the two wavelengths on the same image can be used to estimate the vertical wavelength [Vargas et al., 2007]. An important assumption is that the wavefront shape remains constant between the two images. Our method computes the pattern correlation coefficient (PCC) [Schmetz et al., 1993] between two images, defined by PCCij ðn; mÞ ¼
s12 ðn; mÞ s1 s2 ðn; mÞ
ð1Þ
For each point on the first image, located at coordinates (i, j), a square area of K*K pixels centered on this point is defined, and the same square area is defined on the second image centered on the coordinates (i + n, j + m). s1 is then the standard deviation of the signal over the area of the first image centered on the (i, j) coordinates, and s2(n, m) is the standard deviation of the signal over the area of the second image centered on (i + n, j + m) coordinates. s12(n, m) is the covariance of zero mean signals over the two areas. Values of PCCij(n, m) range from 1 to +1 and are calculated over the range [M2 ; M2 ] of indices n and m. The maximum value = PCCij(n0, m0)) indicates the image of this function (PCCmax ij shifts (n0, m0) at which the two images are best correlated. In order to improve the quality of image shifts retrieved, a minimum value criterion on the correlation is required by > 0.8, and only points with all neighbors fitting PCCmax ij this criterion are retained to avoid image border effects. Additionally, the zero shifts values (n0 = 0 and m0 = 0) are
excluded because instrument effects generally report a signal at zero shifts. Once the best correlated parts in the images are determined and selected, the image shifts are converted into real displacement between the pixels, and then interpreted as wavefront velocities. This image correlation algorithm has been validated by comparison with outputs of COSI-CORR software [Leprince et al., 2007]. Figure 3 shows examples of consecutive images after filtering and projection at Nadir. In Figure 3, the displacements of wavefronts are clearly visible. 3.2. Wave Structure and Distribution [14] The VIRTIS-M data containing wave-like features after processing are summarized in Table 1 with corresponding data parameters. These data have been selected from the large VIRTIS-M database because the same region is mapped twice during the same orbit, enabling the computation of the horizontal displacement of wave features at the same wavelength. The VIRTIS-M data covering the 4.3 mm emission on the dayside at nadir are restricted to south pole or terminator views owing to thermal constraints on the instrument. [15] The two non-LTE emission peaks at the 4.277 mm and 4.315 mm wavelengths show the same patterns after data processing. The image correlation algorithm was not able to detect any coherent horizontal displacement of these patterns between these two wavelengths on the same image. In the rest of the study, only results at the 4.315 mm wavelength are presented because the signal to noise ratio is lower at the 4.277 mm wavelength.
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Table 1. Summary of VIRTIS-M Data Providing More Than 100 Horizontal Velocity Estimates and Corresponding Processing Parameters Orbit
Session
Pixel Size (km)
Range or Dominant Horizontal Wavelength (km)
K/2
M/2
Mean Local Time (h)
219 220 220 232 233 251 473 474 476
04 01 04 01 01 08 00 – 01 00 – 01 00 – 01
16 16 16 16 16 16 16 16 16
340 250 230 210 275 230 90 – 400 110 – 320 90 – 400
10 10 10 10 10 10 12 12 12
10 10 10 10 10 10 12 12 12
11.3 11.5 11.3 11.8 11.8 12.0 12.3 13.5 12.2
[16] The horizontal wavelengths of the upper atmosphere wave features are in the range 90 –400 km. The maximum pixel size of VIRTIS-M data in high space resolution mode (about 16.5 km) should make it possible to resolve horizontal wavelengths below 90 km, but such features are rarely seen. Moreover, the various filtering procedures tested do not reveal horizontal wavelengths larger than 400 km. We therefore believe that the horizontal wavelength range of 90– 400 km is not biased by instrument resolution or data processing. [17] Figure 4 shows an example of the typical distribution of wave features observed from south polar view. For comparison between the three images at the 4.315 mm wavelength, the normalized filtered images are plotted with the same scale (±2.5%). On the composite image, coherent circular wavefronts with horizontal wavelengths larger than 200 km are visible in the middle of the image indicating a possible source region close to the south pole. The wave activity is low on the morningside (on the right). On the eveningside (on the left), small horizontal wavelength features (90 – 200 km) are observed close to dusk with wavefronts almost parallel to the terminator line. The geographical distribution of wave features varies between one orbit and the next, and a wide range of horizontal wavelengths can be observed on the same image, as shown in Figure 2. However, the orientation of wavefronts and the low level of wave activity on the morningside are persistent features. [18] The root-mean-square (RMS) amplitudes of wave features on all images processed are presented in Figure 5 for the two emission peaks at the 4.277 mm and 4.315 mm wavelengths. The areas over which RMS amplitudes are
computed correspond to the image points at which a horizontal velocity vector has been retrieved. The amplitudes of wave-like perturbations of CO2 non-LTE emissions are slightly larger at the 4.315 mm than at 4.277 mm wavelength. The ratio of RMS wave amplitudes is about 3/4. If the emission radiances are directly interpreted in terms of wave amplitudes, this observation may suggest a wave dissipation between the two emission centroids. However, the unknown sensitivities of non-LTE emissions to density perturbations convolute the wave signal preventing the direct interpretation of emission perturbation amplitudes in terms of wave amplitudes. 3.3. Wave Dynamics [19] The horizontal displacement of the wave features is computed for the second peak of CO2 non-LTE emissions (around 4.315 mm) by using the algorithm described in step 5 of data processing. The parameters of the cross-correlation analysis must be adapted to the dominant horizontal wavelength of the signal and to the phase shift between the two images. As shown in Figure 3, the image correlation window size (K) is chosen to be larger than the dominant horizontal wavelength. The displacement range of the correlation window (±M/2) is set to be equal to half the size of the of correlation window (K/2) in order to avoid aliasing. The validity of this parameter was visually checked by comparing image pairs (Figure 3). Parameters K/2 and M/2 chosen for each image pair are given in Table 1, and vary in the range 10– 12 pixels. In addition to the quality tests described in the algorithm, the relative difference between horizontal displacement vectors estimated at 4.277 mm and 4.315 mm is computed. Displacement estimates with a relative difference
Figure 4. Venus south polar views of normalized filtered images (Nij) at the 4.315 mm wavelength on orbit 232. Black and white color scale ranges from 2.5 to 2.5% on the three images. Latitude contours are indicated in degrees, the terminator line is almost horizontal at the lowest border of the images, and the eveningside (morningside) is on the left (right) side. The data processing generates spurious vertical stripes at the limb owing to the strong contrast there. 5 of 11
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Figure 5. Root-mean-square amplitude (in percent) of wave-like features at the 4.315 mm wavelength (circles) and at 4.277 mm (grey dots) on each image processed as a function of average local time (in hours) in reverse scale (eveningside on the left).
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larger than 15% are excluded. From the horizontal displacement estimates and knowing the time lag between the two image points, horizontal phase velocity vectors are calculated. [20] Some examples of velocity vector estimates are shown in Figure 6 for orbits 473 and 474. The wave velocities are obtained in the part of the image that shows wave features with strong contrasts and which maintain the same shape during the time lag between the two images (1800 s). The phase velocity vector is not perpendicular to the wavefront and it presents a strong westward zonal component. [21] The horizontal velocity vectors for the whole data set are plotted in Figure 7. The amplitudes of the phase velocity vectors are in the 50– 110 m/s range. The direction of velocity vectors is zonal westward in the direction of lower atmosphere superrotation, with a small meridional component. The coherency of the velocity vectors is remarkable, because these estimates are obtained from different images along different orbits with different conditions of observation. Meridional and zonal velocities are plotted as a function
Figure 6. Venus south polar views of enhanced wave features at the 4.315 mm wavelength, and corresponding estimates or horizontal wave velocities along orbits (a) 473 and (b) 474. One velocity estimate in 15 is presented to clarify the plots. Local time is converted into longitude (zero longitude is noon) with evening on the left side. 6 of 11
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2.9 mm wavelength. So, the reflected light on cloud top is not the main contribution at the 4.3 mm wavelength. [23] Our data processing scheme is limited to integer pixel displacements, but this limitation will be overcome in future subpixel versions by interpolation of the correlation function or by analysis in the frequency domain [Leprince et al., 2007]. The main limitation comes from the low signal to noise of the wave features partly owing to their low amplitude relative to background emission and partly owing to operational constraints on the spacecraft orientation and on VIRTIS-M exposure time. However, the strong consistency of our results, obtained along different orbits for different observation conditions, excludes the possibility of these wave features having been created by noise or artifacts. [24] Another limitation of our analysis arises from the integrated character of the 4.3 mm radiance along the line of sight. In order to fully interpret these data it is necessary to know the large wavelength space/time variations of nonLTE CO2 volumetric emission, and the sensitivity of these emissions to density perturbations. The first parameter will be obtained from an analysis of the whole VIRTIS data set, including limb profiles, and the second will be modeled as an extension of previous modeling [Rolda´n et al., 2000; Lopez-Puertas and Lopez-Valverde, 1995; Lo´pez-Valverde et al., 2007].
4. Interpretation Figure 7. Estimates of horizontal wave velocities. (top) Norm of the horizontal velocity vector (m/s) as a function of local time (h) in reverse scale (eveningside on the left). (bottom) Norm and direction of horizontal velocity vectors as a function of latitude and local time converted into longitude (zero longitude is noon) with evening on the left side. In this plot, one velocity estimate in 15 is presented to clarify the plots. of local time and latitude on Figure 8. The meridional component is positive (in a northern direction) with an average value of about 30 m/s. The zonal component is westward with an average value of about 70 m/s. The only possible trend in these plots is an increase of the meridional component when approaching dusk, but the low number of images processed in this study did not allow us to validate this feature fully. 3.4. Possible Artifacts and Limitations [22] The radiance observed at nadir is the sum of the direct non-LTE emission and possible other components coming from the reflection of these emissions and/or the sunlight on cloud top at these wavelengths. In order to check the possible contribution of light reflected on cloud top, the wave enhancement procedure was applied on the same images at the 2.9 mm wavelength and in the visible for which the reflected sunlight is the main contribution. Some wave features were detected on these images, but the positions and orientations of these features are completely different from the ones observed at the 4.3 mm wavelength. Similarly, the background emission observed at the 4.3 mm wavelength and its large horizontal wavelength variations are different from the ones observed in the visible and at the
4.1. Wave Type [25] In order to identify the type and structure of the observed wave features, the results are analyzed in the framework of gravity and acoustic wave propagation theories. The difficulty when discussing the wave type (gravity or acoustic waves) lies in the following two major unknowns: the vertical wave structure and the background circulation. The vertical wave structure can be constrained by the thickness of the emission layers and the fact that the two main non-LTE emissions at 4.277 mm and 4.315 mm display the same wave features whereas their centroid altitudes seen at the limb are separated by about 15 km. These two observations impose a minimum vertical wavelength of approximately 30 km. Concerning the background thermospheric circulation, previous models [Zhang et al., 1996; Bougher et al., 1997] predict a strong subsolar to antisolar circulation with winds in the range 60– 130 m/s in this altitude range (110 –140 km) and in this angular distance range to the terminator (15°– 35°). However, the particularities of the atmospheric dynamics in the polar region [Piccioni et al., 2007] can modify this simple picture significantly. [26] The wavefronts are almost parallel to lines of constant latitude, so the horizontal component of the apparent phase velocity (Ch0) can be approximated by the meridional component of the phase velocity which is perpendicular to the wavefronts. An average apparent horizontal phase velocity is Ch0 = 30 ± 10 m/s. If we assume that our wave features are acoustic waves, the intrinsic horizontal phase velocity is Chi = Ch0 U, with the background meridional wind U in the range [130;0] m/s, which is a very conservative hypothesis. We deduce that Chi is in the range [20;170] m/s taking into account error bars. The speed of sound at these altitudes is about 215 m/s. Because this value is a minimum estimate of the intrinsic horizontal phase
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Figure 8. (top) Meridional and (bottom) zonal components of the apparent phase velocities as a function of (left) local time in reverse scale and (right) latitude. velocity Chi for acoustic waves, obtained only for horizontal wave propagation, our wave features cannot be acoustic waves. So, the gravity wave hypothesis is the only remaining possibility. [27] In conclusion, the wave features presented here are indicative of gravity wave phenomena. We will explore in the next two sections the source and propagation of these waves and their links to the background atmospheric dynamics. 4.2. Gravity Wave Source [28] As shown in Figures 4 and 6 the main wavefronts are almost circular and parallel to lines of constant latitude. This observation suggests a source close to the south pole. Figure 9 shows a comparison between the enhanced wave features at the 4.315 mm wavelength and the polar vortex observed at the 5.05 mm wavelength in the same images. This comparison has also been performed on polar vortex features observed at the 3.3 mm wavelength (which is less sensitive to saturation problems) for all the images processed in this study. The direction perpendicular to wavefronts is always pointing in the direction of the region of strongest polar vortex activity, whatever the position and shape of the vortex. These observations clearly identify the polar vortex as the source of the gravity waves observed in the 110–140 km altitude range. This conclusion is also supported by theoretical studies [Baker et al., 1999; Nappo, 2002] which demonstrate that turbulence is able to generate gravity waves propagating upward. For such a source close to south pole, the gravity waves propagate with horizontal components of intrinsic phase and group velocities along meridional direction from the pole to the equator.
[29] Moreover, the analysis of VIRTIS-M images has shown that the wave features at the 4.315 mm wavelength are more numerous with larger amplitude when the polar vortex is on the dayside than on the nightside. As an example, Figures 10a and 10b show that only a few wave features of low amplitude are visible when the polar vortex is completely on the nightside. This observation suggests that the gravity wave propagation from the cloud layer up to the 110 –140 km altitude range is more favorable on the dayside than on the nightside. Figure 10c shows the BruntVa¨isa¨la¨ frequency profile computed for day and night Venus atmosphere parameters [Hunten et al., 1983]. In the 100– 140 km range, it increases on the nightside and decreases on the dayside. The shape of the Brunt-Va¨isa¨la¨ frequency profile on the dayside suggests that gravity waves may be trapped between altitudes of 60 km to 140 km for wave frequencies between 3 and 3.5 mHz. In this layer, trapped waves propagate horizontally with horizontal wavelengths depending on the trapped mode number, and with low attenuation along their path [Lighthill, 2001]. The presence of these trapped waves at frequencies higher than 3 mHz could explain the difference between day and nightside observations. Moreover, the presence of strong background winds and vertical wind variations may also play an important role in the day/night difference of wave propagation by creating wind ducts [Nappo, 2002; Snively et al., 2007; Zhou and Morton, 2007]. 4.3. Implications for Background Circulation [30] The gravity waves at these heights propagate within the global thermosphere circulation. The horizontal phase
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Figure 9. Venus south polar views along orbits (a and b) 220, (c and d) 232, and (e and f) 251 of (left) enhanced wave features at 4.315 mm wavelength and (right) emission at 5.05 mm wavelength which reveals the structure and position of the south polar vortex. Latitude contours are indicated in degrees, and the terminator line is clearly visible through the spurious features created by the data filtering on the nightside. speeds obtained therefore include the effects of both wave propagation and background wind. Consequently, the observed gravity wave propagation can constrain the background wind structure. [31] If we assume that the source of the gravity waves is the turbulent area of the polar vortex close to the south pole, their propagation without wind should be mainly along
meridians. However, a strong zonal component of the horizontal phase velocity is observed in Figures 7 and 8. This observation implies that a zonal westward background wind of about 70 m/s is present at the altitude of 4.315 mm wavelength non-LTE emission perturbations in the 0900 – 1500 LT range in the polar regions. The gravity wave propagation therefore suggests that the zonal lower
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Figure 10. Venus south polar views along orbit 253 of (a) enhanced wave features at the 4.315 mm wavelength and (b) emission at the 5.05 mm wavelength revealing the structure and position of the south polar vortex. Latitude contours are indicated in degrees. (c) Plot of the Brunt-Va¨isa¨la¨ frequency (in mHz) as a function of altitude (in km) on the dayside (solid line) and on the nightside (dotted line), computed for Venus atmosphere models [Hunten et al., 1983]. thermosphere circulation in the 0900 – 1500 LT range is in the direction of the lower atmosphere superrotation up to 30° from the south pole and up to an altitude of at least 115 km. [32] Moreover, because of the absence of phase difference between the wave features observed at the 4.277 mm and 4.315 mm wavelengths, the vertical wavelength of these waves must be larger than twice the centroid altitude difference of these two emissions (15 km at the limb) [Vargas et al., 2007]. In order to use this constraint, the gravity wave dispersion relation with constant background wind should be written [Li et al., 2007] N2 2
ðC0 U Þ
¼ m2 þ k 2 þ
1 4H 2
ð2Þ
where m = 2p lz is the vertical wave number with lz the vertical wavelength, N = 2p f = 0.0215 rad/s (or f = 3.4 mHz) is the Brunt-Va¨isa¨la¨ pulsation which is estimated from a dayside Venus atmosphere model at a latitude of 120 km [Hunten et al., 1983], C0 is the apparent wave velocity, U is the meridional wind, k = 2.513 105 rad/m is the horizontal wave number assuming an average 250 km horizontal wavelength, and H = 4 km is the density scale height at 120 km altitude. If we assume an intrinsic wave propagation from the pole to the equator, and a lower bound on the vertical wavelength lz > 2 * 15 km = 30 km and on the apparent phase velocity C0 > Ch0 = 30 m/s, by using equation (2) we deduce an upper bound on the meridional phase velocity U < 58 m/s. This analysis suggests that the meridional wind is opposite to the wave propagation direction with an amplitude greater than about 60 m/s. This conclusion is in favor of a subsolar to antisolar thermospheric meridional circulation consistent in direction and amplitude with previous results [Zhang et al., 1996; Bougher et al., 1997]. [33] In addition to the gravity wave dispersion analysis, the previous conclusion is also supported by the geographical distribution of gravity wavefronts. Sun et al. [2007]
demonstrated recently that gravity waves propagating against the wind direction are amplified relative to other propagation directions. Assuming a wave source close to the south pole and a strong subsolar to antisolar wind, this prediction is consistent with our observation of gravity wave features above the noise level only in the 0900– 1500 LT range. The wave feature distribution therefore also supports the presence of a strong subsolar to antisolar wind at the south pole.
5. Conclusion [34] The nadir images of non-LTE CO2 emissions present wave-like perturbations of about 0.5% RMS amplitude relative to background signal. These features are mainly observed in the 0900 – 1500 LT range, and their horizontal wavelengths are in the 90 – 400 km range. The dynamics of these wave features are similar for all the images processed along different orbits. Average zonal and meridional components of the horizontal phase velocities are respectively 70 m/s westward and 30 m/s northward. The horizontal phase velocities and horizontal wavelengths are compatible with gravity waves propagating in the 110 – 140 km altitude range which is sampled by non-LTE CO2 emissions at the 4.3 mm wavelength. The considerable uncertainty on the altitude of these perturbations restricts the acquisition of information on the vertical structure of the waves. However, the geographical distribution and the orientation of wavefronts clearly identify the polar vortex at work in the cloud layer as the source of gravity waves. Finally, the characteristics of gravity wave propagation can be used to infer the background atmospheric winds. The strong zonal component of gravity wave phase velocities is not due to their propagation from a source close to the south pole, and can consequently be attributed to the background atmospheric circulation in the 110 – 140 km altitude range. The gravity wave dispersion relation and the geographical distribution of wavefronts strongly argue in favor of the presence of a subsolar to antisolar wind of amplitude larger than 60 m/s close to the south pole.
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[35] This first direct mapping of thermospheric gravity waves demonstrates the strong influence of the polar vortex on the atmospheric polar circulation of Venus [Piccioni et al., 2007], and illustrates an upward energy transfer from the cloud layer to the thermosphere. Moreover, this study seems to demonstrate the persistence of superrotation dynamics at least up to 115 km in polar regions (up to 30° from the south pole). Furthermore, our data support the existence of the subsolar to antisolar circulation predicted by thermosphere circulation models [Zhang et al., 1996; Bougher et al., 1997], with meridional wind speeds greater than 60 m/s in this polar region. [36] The increase of the VIRTIS database will improve the statistics of wave features, and its combination with VMC data [Markiewicz et al., 2007] will make it possible to follow gravity wave propagation from the top of the cloud layer up to the thermosphere. In order to improve our description and understanding of these wave phenomena, a large-scale threedimensional model of CO2 volumetric emission rates should be developed by using the limb and nadir VIRTIS observations, the sensitivity of these emissions to gravity wave perturbations should be modeled, and the gravity wave propagation within the background circulation should be computed. Finally, these high-altitude emissions can also be used to investigate transient infrasonic signals generated by quakes and volcanic events [Garcia et al., 2005]. [37] Acknowledgments. Two anonymous reviewers improved this paper considerably by their constructive comments. We thank Ste´phane Erard and Alejandro Cardesin for computation and validation of VIRTIS-M data and corresponding observation geometry, Ricardo Hueso Alonso for his help with IDL data projection routines, F. Ayoub for his help with COSI-CORR software, Se´bastien Lebonnois for helpful discussions on the atmosphere dynamics, and P.E. Mallet for validating the correlation software. This study was funded by CNES through space research scientific projects.
References Alexander, M. (1992), A mechanism for the Venus thermospheric superrotation, Geophys. Res. Lett., 19, 2207 – 2210. Baker, R., G. Schubert, and P. Jones (1999), High rayleigh number compressible convection in Venus’ atmosphere: Penetration, entrainment, and turbulence, J. Geophys. Res., 104, 3815 – 3832. Bougher, S., and W. Borucki (1994), Venus O2 visible and IR nightglow: Implications for lower thermosphere dynamics and chemistry, J. Geophys. Res., 99, 3759 – 3776. Bougher, S., M. Alexander, and H. Mayr (1997), Upper atmosphere dynamics: Global circulation and gravity waves, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher et al., pp. 259 – 291, Univ. of Ariz. Press, Tucson. Drossart, P., et al. (2004), VIRTIS imaging spectrometer for the ESA/Venus Express mission, in Earth Observing Systems IX, vol. 5583, pp. 175 – 185, Int. Soc. for Opt. Eng., Bellingham, Wash. Drossart, P., et al. (2007), A dynamic upper atmosphere of Venus as revealed by VIRTIS on Venus Express, Nature, 450, 641 – 645, doi:10.1038/nature06140. Garcia, R., P. Lognonne´, and X. Bonnin (2005), Detecting atmospheric perturbations produced by Venus quakes, Geophys. Res. Lett., 32, L16205, doi:10.1029/2005GL023558. Gierasch, P. (1987), Waves in the atmosphere of Venus, Nature, 328, 510 – 512. Gierasch, P., et al. (1997), The general circulation of the Venus atmosphere: An assessment, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher et al., pp. 459 – 500, Univ. of Ariz. Press, Tucson. Gilli, G., M. Lo´pez-Valverde, P. Drossart, G. Piccioni, S. Erard, and A. Cardesı´n-Moinelo (2009), Limb observations of CO2 and CO non-LTE emissions in the Venus atmosphere by VIRTIS/Venus Express, J. Geophys. Res., doi:10.1029/2008JE003112, in press. Hinson, D., and J. Jenkins (1995), Magellan radio occultations measurements of atmospheric waves on Venus, Icarus, 114, 310 – 327. Hunten, D., L. Colin, T. Donahue, and V. Moroz (Eds.) (1983), Venus, 1143 pp., Univ. of Ariz. Press, Tucson.
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Kasprzak, W. T., H. B. Niemann, A. E. Hedin, and S. W. Bougher (1993), Wave-like perturbations observed at low altitudes by the Pioneer Venus Orbiter Neutral Mass Spectrometer during orbiter entry, Geophys. Res. Lett., 20, 2755 – 2758. Leprince, S., S. Barbot, F. Ayoub, and J. Avouac (2007), Automatic and precise ortho-rectification, coregistration, and subpixel correlation of satellite images: Application to ground deformation measurements, IEEE Trans. Geosci. Remote Sens., 45, 1529 – 1558. Li, F., G. Swenson, A. Liu, M. Taylor, and Y. Zhao (2007), Investigation of a wall wave event, J. Geophys. Res., 112, D04104, doi:10.1029/ 2006JD007213. Lighthill, J. (2001), Waves in Fluids, 520 pp., Cambridge Univ. Press, Cambridge, U.K. Lopez-Puertas, M., and M. A. Lopez-Valverde (1995), Radiative energy balance of CO2 non-LTE infrared emissions in the Martian atmosphere, Icarus, 114, 113 – 129, doi:10.1006/icar.1995.1047. Lopez-Puertas, M., M. Lopez-Valverde, R. Garcia, and R. G. Roble (2000), A review of CO2 and CO abundances in the middle atmosphere, in Atmospheric Science Across the Stratopause, Geophys. Monogr. Ser., vol. 123, edited by D. E. Siskind, S. D. Eckermann, and M. E. Summers, pp. 83 – 100, AGU, Washington D.C. Lo´pez-Valverde, M.A., P. Drossart, R. Carlson, R. Mehlman, and M. RoosSerote (2007), Non-LTE infrared observations at Venus: From NIMS/ Galileo to VIRTIS/Venus Express, Planet. Space Sci., 55, 1757 – 1771, doi:10.1016/j.pss.2007.01.008. Markiewicz, W., et al. (2007), Morphology and dynamics of the upper cloud layer of Venus, Nature, 450, 633 – 636, doi:10.1038/nature06320. Mayr, H. G., I. Harris, W. T. Kasprzak, M. Dube, and F. Varosi (1988), Gravity waves in the upper atmosphere of Venus, J. Geophys. Res., 93, 11,247 – 11,262. Nappo, C. (2002), An Introduction to Atmospheric Gravity Waves, 276 pp., Academic, San Diego. Piccioni, G., et al. (2007), South-polar features on Venus similar to those near the north pole, Nature, 450, 637 – 640, doi:10.1038/nature06209. Rolda´n, C., M. Lo´pez-Valverde, and M. Lo´pez-Puertas (2000), Non-LTE infrared emissions of CO2 in the atmosphere of Venus, Icarus, 147, 11 – 25. Schmetz, J., K. Holmlund, J. Hoffman, B. Strauss, B. Mason, V. Gaertner, A. Koch, and L. Van De Berg (1993), Operational cloud-motion winds from Metosat infrared images, J. Appl. Meteorol., 32, 1206 – 1225. Snively, J. B., V. P. Pasko, M. J. Taylor, and W. K. Hocking (2007), Doppler ducting of short-period gravity waves by midlatitude tidal wind structure, J. Geophys. Res., 112, A03304, doi:10.1029/2006JA011895. Sun, L., W. Wan, F. Ding, and T. Mao (2007), Gravity wave propagation in the realistic atmosphere based on a three-dimensional transfer function model, Ann. Geophys., 25, 1979 – 1986. Vargas, F., G. Swenson, A. Liu, and D. Gobbi (2007), O(1S), OH, and O2(b) airglow layer perturbations due to AGWs and their implied effects on the atmosphere, J. Geophys. Res., 112, D14102, doi:10.1029/ 2006JD007642. Woo, R., J. W. Armstrong, and A. Ishimaru (1980), Radio occultation measurements of turbulence in the Venus atmosphere by Pioneer Venus, J. Geophys. Res., 85, 8031 – 8038. Woo, R., J. W. Armstrong, and A. J. Kliore (1982), Small-scale turbulence in the atmosphere of Venus, Icarus, 52, 335 – 345, doi:10.1016/00191035(82)90116-6. Young, R. E., R. L. Walterscheid, G. Schubert, A. Seiff, V. M. Linkin, and A. N. Lipatov (1987), Characteristics of gravity waves generated by surface topography on Venus: Comparison with the VEGA balloon results, J. Atmos. Sci., 44, 2628 – 2639. Zhang, S., S. Bougher, and M. Alexander (1996), The impact of gravity waves on the Venus thermosphere and O2 IR nightglow, J. Geophys. Res., 101, 23,195 – 23,205. Zhou, Q., and Y. Morton (2007), Gravity wave propagation in a nonisothermal atmosphere with height varying background wind, Geophys. Res. Lett., 34, L23803, doi:10.1029/2007GL031061.
P. Drossart, LESIA, Observatoire de Paris, UPMC, Universite´ ParisDiderot, CNRS, 5 place Jules Janssen, F-92195 Meudon, France. R. F. Garcia, Laboratoire de Dynamique Terrestre et Plane´taire, UMR5562, Observatoire Midi-Pyre´ne´es, Universite´ de Toulouse, 14 avenue Edouard Belin, F-31400 Toulouse, France. (
[email protected]) M. Lo´pez-Valverde, Instituto de Astrofı´sica de Andalucı´a, CISC, 50 Camino Bajo de Hue´tor, ES-18080 Granada, Spain. G. Occhipinti, Institut de Physique du Globe de Paris, Universite´ Denis Diderot Paris 7, 4 place Jussieu, F-75252 Paris CEDEX 05, France. G. Piccioni, INAF, IASF, via del fosso del cavaliere 100, I-00133 Rome, Italy.
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Characterization of mesoscale gravity waves in the upper and lower clouds of Venus from VEX-VIRTIS images J. Peralta,1 R. Hueso,1 A. Sa´nchez-Lavega,1 G. Piccioni,2 O. Lanciano,2 and P. Drossart3 Received 9 May 2008; revised 11 August 2008; accepted 22 August 2008; published 11 December 2008.
[1] Images obtained from the Visible and InfraRed Thermal Imaging Spectrometer
(VIRTIS)-M instrument onboard Venus Express present visible trains of alternating bands of cloud brightness in two different layers: at the upper cloud tops (66 km altitude) observed in the dayside hemisphere using reflected ultraviolet light (380 nm) and in the lower cloud (47 km altitude) observed in the nightside hemisphere using thermal radiation (1.74 mm). The waves are nearly zonal (with the bands perpendicular to latitude circles), have wavelengths of 60–150 km, propagate westward with low phase velocities relative to the zonal flow, and are confined in wave packets of 400 to 1800 km in length. The waves in the lower cloud observed in the infrared are widely distributed around the planet, and their appearance varies widely throughout the VIRTIS data set. The locations of both types of waves seem not correlated with latitude, local times, surface topography, or the structure of the wind. In both cases the characteristics of the waves correspond to gravity waves propagating in confined stable layers of the atmosphere. We examine the properties of these waves in terms of a linear model and perform a simple analysis to discuss the vertical stability of the atmosphere within Venus clouds. Citation: Peralta, J., R. Hueso, A. Sa´nchez-Lavega, G. Piccioni, O. Lanciano, and P. Drossart (2008), Characterization of mesoscale gravity waves in the upper and lower clouds of Venus from VEX-VIRTIS images, J. Geophys. Res., 113, E00B18, doi:10.1029/2008JE003185.
1. Introduction [2] Atmospheric internal gravity waves are wave disturbances whose restoration force is buoyancy under stable stratification of the atmosphere [Holton, 1992]. They are commonly observed in the stratosphere of most planets as oscillations on the temperature field and they have also been observed as wavy structures in cloud fields of the Earth [Houze, 1993], Jupiter [Hunt and Muller, 1979; Flasar and Gierasch, 1986; Reuter et al., 2007] and Venus [Gierasch et al., 1997; Markiewicz et al., 2007]. Since they are supported by such a fundamental force (buoyancy) they are of great interest to atmospheric dynamics. On the one hand, from their properties one can infer the static stability of the atmosphere (since these waves can only propagate in regions where the static stability is positive). On the other, they are responsible of a number of important dynamical phenomena including the transfer of energy and momentum vertically. They are of particular importance to the atmosphere of Venus, where their role as a mechanism able to transport momentum from the surface to the atmospheric 1 Departamento de Fı´sica Aplicada I, Escuela Te´cnica Superior de Ingenierı´os, Universidad del Paı´s Vasco, Bilbao, Spain. 2 Istituto di Astrofisica Spaziale e Fisica Cosmica, Istituto Nazionale di Astrofisica, Rome, Italy. 3 Laboratoire d’Etudes Spatiales et d’Instrumentation en Astrophysique, Observatoire de Paris, UPMC, Universite´ Paris-Diderot, CNRS, Meudon, France.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003185$09.00
mean circulation has been examined by several teams, although without a final consensus on their role in powering the atmospheric superrotation [Hou and Farrell, 1987; Gierasch, 1987; Leroy and Ingersoll, 1995]. Space probes determined during their descent at different latitudes and local times the existence of at least two stable layers: between 30 and 48 km, and upward of 55 km [Kliore and Patel, 1980; Seiff et al., 1985; Gierasch, 1987]. These altitude levels are home to dense clouds on Venus [Esposito et al., 1997] and gravity waves arising on them can manifest as regular patterns on the clouds albedo fields [Markiewicz et al., 2007; McGouldrick and Toon, 2008]. The altitudes below 30 km and the intermediate layer between 48 and 55 km have small static stabilities and are not suitable candidates for the development of gravity waves. [3] Gravity waves observed at the upper cloud level were first inferred from analysis of Mariner 10 [Belton et al., 1976] and Pioneer Venus [Rossow et al., 1980] ultraviolet images of the upper cloud field at z 65 km. More recently, the Venus Monitoring Camera (VMC) onboard Venus Express has also detected regular cloud patterns identified as gravity waves from their observed characteristics [Markiewicz et al., 2007]. Mariner 10 images of equatorial elongated cloud patterns of about 5000 km in the longitudinal direction and separated 500 km in latitude (circumequatorial belts) were interpreted as evidence of internal gravity waves. These structures propagated southward near the equator with phase speeds of 20 m/s [Belton et al., 1976]. The same circumequatorial belts were seen during the Pioneer Venus mission in addition
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to a new type of train of wavelike features composed of dark streaks of 2000 km, separated by 200 km, but in this case lying at a large angle to the constant latitude circles [Rossow et al., 1980]. No evidence for small-scale internal gravity waves or for circumequatorial belts was found in images taken by the Galileo spacecraft in 1990 [Belton et al., 1991; Toigo et al., 1994; Peralta et al., 2007a]. This was linked to a lower abundance of convective clouds visible during the Galileo flyby. The high-resolution images obtained recently by the VMC instrument onboard Venus Express display vigorous convection in near equatorial latitudes and visible wave activity (wavelengths of just a few tens of kilometers) in tropical latitudes in the south hemisphere and in middle to subpolar latitudes in the north hemisphere [Markiewicz et al., 2007]. [4] Evidence for gravity waves in the Venus atmosphere does not come only from imagery. The four Pioneer Venus probes, as they descended through the atmosphere, obtained altitude profiles of temperature (upward of 110 km) and wind velocity (lower cloud region) that showed wavelike variations with vertical wavelengths of 5 – 10 km [Seiff at al., 1980; Counselman et al., 1980]. Similar vertical variations can be seen in infrared temperature soundings [Taylor et al., 1980], and in radio occultation temperatures from Pioneer Venus [Kliore and Patel, 1980], Venera 9 [Kolosov et al., 1980] and Magellan [Hinson and Jenkins, 1995] sensitive to mesospheric vertical levels between the cloud tops (z 65 km) and the thermosphere base (z 100 km). Above these levels, gravity waves with wavelengths between 100 and 600 km have also been detected on the thermosphere as density perturbations studied by the mass spectrometer on the Pioneer Venus orbiter [Niemann et al., 1980; Kasprzak et al., 1988], and from non LTE CO2 emissions observed by the Visible and InfraRed Thermal Imaging Spectrometer (VIRTIS) instrument on Venus Express (R. Garcı´a et al., Gravity waves in Venus upper atmosphere revealed by CO2 non LTE emissions, submitted to Journal of Geophysical Research, 2008). [5] Although several candidates could act as sources of gravity waves (flow over a mountain range, KelvinHelmholtz instability around a jet stream), convective processes generated in the low-stability layers between 48 and 55 km and below 30 km are considered to play the main role for gravity wave generation in the middle and upper regions of the clouds layer [Schubert, 1983; Gierasch, 1987; Leroy and Ingersoll, 1995, 1996; Baker et al., 2000a, 2000b]. [6] In this paper we present the observed characteristics of wavelike features present in the atmosphere of Venus at two cloud levels detected on images at different wavelengths obtained by the VIRTIS instrument [Drossart et al., 2007] onboard Venus Express [Svedhem et al., 2007]. We identify these features as gravity waves and present a simple theoretical approach to interpret them in terms of a ‘‘textbook’’ classical linear model. This is the first time that the abundant and variable mesoscale gravity waves within the lower cloud level of Venus have been characterized. From the comparison between the observations and model we infer the existence of a widely (from latitudes 40° to 80°S) and temporarily extended vertical stable region between 30 and 48 km, in agreement with
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previous inferences [Gierasch, 1987; Leroy and Ingersoll, 1995].
2. Observations [7] The instrument VIRTIS on board Venus Express is an imaging spectrometer (VIRTIS-M) combined with a highresolution IR-spectrometer (VIRTIS-H). VIRTIS-M is capable of obtaining images over several visible and infrared wavelengths in two channels: visible (0.3 to 1 mm) and near infrared (1 to 5 mm) [Drossart et al., 2007]. Dayside observations in the ultraviolet show cloud features in the upper cloud (380 nm), corresponding to an altitude of (66 ± 4) km. At the near infrared 980 nm, solar photons are more penetrating and reach the base of the upper cloud at (61 ± 3) km [Belton et al., 1991; Peralta et al., 2007b]. Night side images of Venus at 1.74 mm are sensitive to the opacity of the lower cloud layer at 44– 48 km, with the radiance coming from altitude levels beneath the clouds and being attenuated as it passes through the lower cloud layer [Carlson et al., 1991; Crisp et al., 1991]. Sa´nchez-Lavega et al. [2008] present a detailed account of the global cloud dynamics observed in these filters and the relation between wavelength and cloud altitude determination. [8] Because of the highly elliptical orbit of the VEX spacecraft, with nearly 80°N pericenter, only the south hemisphere can be covered by the VIRTIS instrument. The spatial resolution of VIRTIS images vary from 15 km/pix at polar and subpolar latitudes to 45 km/pix at equatorial latitudes. The nightside images used in this study cover the period from 12 April 2006 (Orbit Insertion) to 9 March 2007. The dayside images cover the period from Orbit Insertion to 28 July 2007, extending during a large fraction of the nominal mission. The examined period for the visible images was extended compared to the infrared images because of the lower number of orbits containing high-resolution well-contrasted images of the dayside. From these periods, and attending to image quality and spatial resolution, we selected observations corresponding to 112 orbits in the visible and 116 orbits in the infrared. [9] Figure 1 shows maps of the number of observations for each latitude and local time with a high spatial resolution suitable for gravity waves search (spatial resolution better than 30 km/pix). This roughly corresponds to latitudes poleward of 30°S in both day and nightside. [10] The images were navigated, corrected from defects (mainly striations and hot and dark lines), processed and mapped into cylindrical and polar projections using the ‘‘Planetary Laboratory Image Analysis’’ (PLIA) software (R Hueso et al., The Planetary Laboratory for Image Analysis (PLIA), submitted to Planetary and Space Science, 2008.). Each image was processed individually by a combination of contrast enhancement, unsharp mask and smoothing filters.
3. Results [11] Examples of systems of waves observed in the cloud fields in ultraviolet (380 nm), near-infrared (980 nm) and infrared (1.74 mm) wavelengths (hereafter, UV, NIR and IR,
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Figure 1. Global resolution maps and frequency of high-resolution observations made with Visible and InfraRed Thermal Imaging Spectrometer (VIRTIS) on the (a) dayside and (b) nightside of Venus. The number of observations of a specific latitude-local time with a high spatial resolution of 30 km/pix is indicated. This spatial resolution is suitable for observations of waves with wavelengths larger than 100 km. Observations with better resolutions of the order of 15 –20 km/pix were also present in the data set with much less frequency. The position of the waves encountered is shown as symbols (triangles for the waves in the upper cloud, and crosses for the waves in the lower cloud) and correlates with regions with high number of observations at this resolution. respectively) are displayed in Figure 2. The spatially distributed regular patterns of alternative bright and dark stripes represent reflectivity variations in the ultraviolet and near infrared (observed during dayside) and opacity variations in the infrared (observed during nightside). The images show trains of crests and troughs generally ordered in longitude (oriented in the zonal direction) and extending a few degrees in latitude nearly perpendicularly to the wave packet. In many cases the troughs are wider than the crests, or they simply disappear. Some of the wave packets are very regular while others have nonequidistant crests, change their latitudinal dimensions over the packet or show a slight curvature. Beside on Venus, these irregularities have also been observed in mesoscale cloud waves observed in Jupiter and interpreted in terms of ducted gravity waves [Flasar and Gierasch, 1986]. [12] We performed a systematic search of waves (see Figure 1) as those shown on Figure 2, finding 6 distinctive and separate wave packets in ultraviolet images, one wave packet in the near infrared (this wave system was observed simultaneously in the ultraviolet and the infrared with nearly identical characteristics) and 30 wave packets in infrared images. In some cases, a single infrared image presents wave systems at various latitudes. The larger abundance of waves in the infrared images results in a larger diversity of wave properties. The average image contrast variations that characterize these wave systems are on the order of 1% in the ultraviolet, less than 1% in the near infrared and vary widely in the infrared with some faint wavy structures with contrast variations of 1%, subtle to detect, to clearly observable systems with contrast of 15%. This study does not rule out the possible existence of less conspicuous wave variations in the cloud fields with less contrast or shorter wavelengths as found in the upper cloud by the VMC instrument [Markiewicz et al., 2007]. [13] We measured the observable properties of the wave packets such as their latitude and local time location, packet length (i.e., their extension perpendicular to the wavefronts),
packed width (i.e., their extension in the wavefront direction) and orientation respect to the parallels (lines of constant latitude). Results are listed in Tables 1a–1c. [14] Figure 3 shows the location of the clouds in terms of latitudes, local times and surface topography. These maps show abundant waves at the lower cloud with no correlation in their position with surface topography (Figure 3a) or local time (Figure 3b). This is in contrast to the northern hemisphere, where the Vega balloon registered the influence of the surface in atmospheric motions with oscillations at the lower cloud level interpreted as gravity waves generated by the flow over specific mountainous regions [Sagdeev et al., 1986; Young et al., 1987, 1994]. In the southern hemisphere, where the surface is nearly flat compared to the northern hemisphere, especially in the sampled latitudes, the waves seem to be present all along the nightside hemisphere except inside the South Polar Dipole [Piccioni et al., 2007]. This could be due to an observational effect since the dipole interior is particularly dark on the infrared images. In any case, gravity waves have been detected up to the dipole’s outer edge (80°S in orbit 84) and the global dynamics of the polar area will be presented elsewhere. [15] The Venus Monitoring Camera on VEX has found abundant cloud waves of higher frequency (wavelengths of a few tens of km) in the upper clouds from south tropical to north polar latitudes [Markiewicz et al., 2007]. These waves, if also present in the south hemisphere, cannot be resolved by VIRTIS observations. The scarcity of large waves detected on the upper clouds does not allow extracting a clear conclusion about their distribution but they seem clustered in the early morning hours and late afternoon. In both cloud layers, we do not find wave packets equatorward of 40°S. This is probably caused by the comparatively low number of high-resolution observations of those latitudes. However, we would also like to point out that at those latitudes there seems to be a transition to different cloud structures, with dark midlatitude bands in the upper cloud (UV), and a mottled turbulent morphology in the lower
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Figure 2. Examples of cloud waves observed on VIRTIS images projected into polar maps for size comparison. Figures 2a and 2b show waves observed at 380 nm in the upper cloud level. The singular detection of a wave at 980 nm images is shown in Figure 2c and correlates with the same structure observed in 380 nm and shown in Figure 2b. Either this is a vertically extended wave, or the observed levels in those wavelengths can be closer than expected at specific latitudes. Large wave packets observed in 1.74 mm in the lower cloud are shown in Figures 2d and 2e. cloud (IR) that on the one hand may difficult the detection of gravity waves but could also represent a region of active cellular convection where gravity waves might not be generated. [16] The extension covered by the wave packets, i.e., the packet length, varies between 640 and 1200 km in the upper cloud (UV) with a mean value of 920 km and from 280 to 1700 km in the lower cloud (IR) with a mean value of 760 km. Their average width are 320 and 230 km in the in
the upper (UV) and lower cloud (IR), respectively. We further measured the wavelength of each packet, the number of crests and, when possible, their zonal motions. Since the winds at cloud level derived from VIRTIS are variable in terms of the local time [Sa´nchez-Lavega et al., 2008] we also measured the local velocity at nearby longitudes and latitudes of each individual wave packet to better estimate their phase speed relative to the local wind. This was done only on those orbits where we could observe the same cloud
Table 1a. Summary of Wave Packets Properties From UV Images by VEX-VIRTISa Orbit
Date (mm/dd/yy)
Latitude (deg)
Local Hour
Packet Length (km)
Packet Width (km)
Orientation (deg)
59 59 170 255 388
06/18/2006 06/18/2006 10/07/2006 12/31/2006 05/13/2007
41 58 62 70 74
16,4 15,7 06,8 15,0 08,1
>860 >865 1275 980 640
170 340 340 420 335
25 30 25 40 2
a
Visible and InfraRed Thermal Imaging Spectrometer (VIRTIS).
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Table 1b. Summary of Wave Packets Properties From NIR Images by VEX-VIRTIS
Orbit
Date (mm/dd/yy)
Latitude (deg)
Local Hour
Packet Length (km)
Packet Width (km)
Orientation (deg)
59
06/18/2006
58
15.4
>860
350
35
field with time differences large enough to track the motions (1 h). Repeated observations of the same wave features indicate they conserve their morphology and albedo variations over these timescales being nondispersive. These properties appear detailed in Tables 2a– 2c. [17] Observed wavelengths are within the range of 90 to 170 km in the upper cloud (UV) and 60 to 180 km in the lower cloud but we cannot rule out waves of higher frequencies (wavelengths shorter than 50 km) not resolved by our observations. At both cloud levels the phase velocities turn out to be generally low respect to the averaged zonal flow. This is in agreement with our interpretation in terms of gravity waves that will be later discussed. [18] The overall properties of these wave systems do not seem to vary with latitude or local time. The wavelength behavior and wave packet properties in terms of latitude appear on Figure 4. Neither the wavelength (Figure 4a), nor the wave packet length (Figure 4b) or the wave packet width (Figure 4b) seem to be correlated with latitude. Only the wave packet orientation (Figure 4d) seems to slightly grow at higher (more poleward) latitudes. The wave system detected on the outer edge of the dipole is aligned with this
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structure and presents the largest inclination with respect to the parallels. [19] The derivation of the properties detailed above critically depends on the quality of the data and, for the determination of the wavelengths and phase speeds, in the ability to accurately track specific structures in the wavefronts and outside them. We estimate the navigation error to be lower than 1 pixel but since all the measurements were done by a human operator the measurement error can be larger. We therefore estimate the wavelengths to be accurate to a 10%, packet lengths and widths accurate to a few tens of km and phase speeds to have a mean measurement error of 8 m/s. The sometimes subtle identification of wave patterns in ultraviolet images relied on the confirmation of all wave structures by a second operator.
4. Discussion [20] On the basis of the characteristics of the wave systems described above we interpret them as internal gravity waves lying at regions of positive static stability. On the one hand, the high frequency of the waves and the slow rotation of the planet discard the observed waves as manifestations of Rossby waves and on the other, these frequencies are much smaller than the acoustic cut-off frequency. Also their latitudes and short sizes discard Kelvin waves as the underlying mechanism for the regular cloud patterns. Roll convection on Earth sometimes produces structures similar in their shape to the regular alignment of clouds observed on our images but this phenomenon is characteristic of the planetary boundary layer and depends on a critical equilibrium between a small static stability and
Table 1c. Summary of Wave Packets Properties From IR Images by VEX-VIRTIS Orbit
Date (mm/dd/yy)
Latitude (deg)
Local Hour
Packet Length (km)
Packet Width (km)
Orientation (deg)
84 96 97 97 100 112 112 112 112 112 113 114 118 141 142 161 164 166 228 231 261 290 313 315 317 317 317 318 321 323
07/13/2006 07/25/2006 07/26/2006 07/26/2006 07/29/2006 08/10/2006 08/10/2006 08/10/2006 08/10/2006 08/10/2006 08/11/2006 08/12/2006 08/16/2006 09/08/2006 09/09/2006 09/28/2006 10/01/2006 10/03/2006 12/04/2006 12/07/2006 01/06/2007 02/04/2007 02/27/2007 03/01/2007 03/04/2007 03/04/2007 03/04/2007 03/05/2007 03/07/2007 03/10/2007
80 48 39 61 55 60 49 49 53 58 54 54 65 65 64 58 50 53 51 78 59 63 60 61 53 59 63 60 49 41
21,9 02,0 22,1 19,8 03,5 23,8 23,1 22,3 21,4 21,1 01,5 05,1 00,2 20,5 03,8 03,6 03,6 03,2 00,9 03,3 21,8 20,6 02,8 01,2 19,1 19,7 20,5 19,5 19,7 19,4
>600 1750 660 570 >1525 610 410 730 785 575 >1140 >645 635 520 >1415 1050 335 >790 690 280 620 >440 >860 760 >1230 685 450 640 >1300 615
225 520 160 160 270 105 170 320 200 205 180 205 265 370 160 310 380 225 370 130 240 250 315 335 170 200 145 150 210 200
50 15 35 40 2 20 40 25 8 7 17 14 2 11 13 17 20 12 10 30 13 4 11 0 10 15 7 0 12 0
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Table 2a. Summary of Wave Properties From UV Images by VEX-VIRTIS Orbit 59 59 170 255 388
Latitude Number of Wave Length Crest Width cx (deg) Crests (km) (km) (m/s) 41 58 62 70 74
8 7 5 11 4
95 115 210 90 170
45 60 125 45 100
113 108 62 29 76
j~ cju (m/s) 12 40 0.5 14 36
a dynamic instability due to the vertical wind shear (low values of the Richardson number). Our wave systems are located several tens of km above the surface and at regions of high Richardson numbers. [21] In order to gain insight into the nature of the waves and further determine their characteristics, we will examine their wave dispersion relation using a simple ‘‘textbook’’ classical linear model of gravity waves. The wave dispersion relation can be deduced by neglecting the effects of rotation under the Boussinesq approximation and assuming two-dimensional motion. Linear perturbations over a basic state with constant zonal flow and density yield the dispersion relation [Holton, 1992]: ðw uk Þ2 ¼
k2
k2 N 2; þ m2
ð1Þ
is the where w is the frequency of the wave perturbation, u mean wind projected in the direction of the wave phase propagation, N2 is the square of the buoyancy or BruntVa¨isa¨lla¨ frequency, and k and m are the horizontal and vertical wave numbers related to the horizontal and vertical wavelengths, lx and lz, by k = 2p/lx and m = 2p/lz. From (1) it is straightforward to deduce the horizontal and vertical components of the phase speed vector relative to the mean flow as,
Figure 3. Polar map of cloud waves’ locations in terms of their longitude, local time, morphology of the clouds, and surface topography. Figure 3a shows the wave packets central positions (triangles for the waves in the upper cloud, and crosses for the waves in the lower cloud) over a surface map obtained by the Magellan IDRS instrument [Saunders et al., 1991]. The topography varies from 1.3 km in the darkest low plains to +2.5 km at the top of the highest mountains in this projection. The higher concentration of waves in the sector between longitudes 210 – 300° is due to a bias in the spatial sampling. Figure 3b shows the cloud waves in terms of latitude and local time displaying the typical cloud structure at these levels with a background image from the Venus Orbit Insertion covering the whole southern hemisphere. Given the spatial resolution and sampling of the polar region, the lack of cloud waves at these latitudes might be significant.
N cx ¼ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 2 k þ m2
ð2aÞ
N k cz ¼ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi ; k 2 þ m2 m
ð2bÞ
where cx is the phase speed relative to the mean wind flow. We now proceed to verify if those waves for which we have information about their phase speeds verify this simple model of internal gravity waves. [22] The clouds observed in the ultraviolet are located at 66 km altitude. These levels are close to the transition between the troposphere and the mesosphere, where the static stability of the atmosphere abruptly increases. A reasonable value of the Brunt-Va¨isa¨la¨ frequency from
Table 2b. Summary of Wave Properties From NIR Images by VEX-VIRTIS Orbit 59
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Latitude Number of Wave Length Crest Width cx j~ cju (deg) Crests (km) (km) (m/s) (m/s) 58
7
125
75
-
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Table 2c. Summary of Wave Properties From IR Images by VEX-VIRTIS
Orbit
Latitude (deg)
Number of Crests
Wave Length (km)
Crest Width (km)
cx (m/s)
j~ cju (m/s)
84 96 97 97 100 112 112 112 112 112 113 114 118 141 142 161 164 166 228 231 261 290 313 315 317 317 317 318 321 323
80 48 39 61 55 60 49 49 53 58 54 54 65 65 64 58 50 53 51 78 59 63 60 61 53 59 63 60 49 41
6 12 8 8 12 8 5 7 6 4 6 5 7 4 18 14 4 6 5 5 4 4 7 6 9 9 6 5 12 5
63 155 100 80 130 60 103 90 125 105 140 100 90 180 75 75 100 130 130 105 145 120 135 100 100 90 90 100 105 110
40 70 35 30 65 30 50 20 70 45 90 40 50 75 44 40 50 60 65 65 65 66 65 55 55 50 45 50 50 55
26 64 59 58 44 61 64 45 42 60 64 59 60 66 54 50
1.0 9.0 3.1 2.7 0.2 5.1 1.8 29.7 9.0 5.0 3.1 2.4 1.1 6.1 7.5 3.3
different probes for these levels is N2 260 106 s2 [Seiff et al., 1985; Del Genio and Rossow, 1990; Gierasch et al., 1997]. Above these levels the stability of the atmosphere increases to higher values characteristic of the mesosphere. Below the clouds the stability decreases to nearly zero allowing for a region between 48 and 55 km where vertical convection can develop. The lower clouds acting as a source of opacity for the atmospheric emitted radiation are located just below, in altitudes (43 – 48 km) characterized by high values of the static stability. We consider as a representative value for the Brunt-Va¨isa¨la¨ frequency in the lower cloud N2 65 106 s2. This is close to the peak values measured by the Pioneer Venus Night Sounder at the height of 45 km [Seiff et al., 1985; Gierasch et al., 1997] over the stable layer situated in the height range 30– 48 km, and it is within the range of 50– 100 106 s2 obtained at these altitudes by the ensemble of Pioneer Venus sounders. This stable layer may act as a duct where waves can form excited by convection from the lower levels or from downdrafts from the unstable levels above [Baker et al., 1998]. [23] Figure 5 shows the behavior of the gravity wave dispersion relationship for different values of the vertical wavelength lz and the assumed values of the Brunt-Va¨isa¨la¨ frequency. The upper cloud waves (Figure 5a) seem to be distributed in two groups of short (5 km) and large (15 km) vertical wavelengths or in regions with small and large values of the atmospheric static stability. We note that the cloud wave simultaneously detected in the ultraviolet and
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near infrared images corresponds to one of the two cases of long vertical wavelengths. The lower cloud waves (Figure 5b) are more abundant and are homogeneously distributed in the plot in regions of short vertical wavelengths (5 – 10 km). This is coherent with the vertical extent of the duct and consistent with the fact that more vertically extended waves would be dispersed by the interaction with the vertical wind shear [Leroy and Ingersoll, 1995]. [24] In both cases, phase velocities relative to the mean zonal flow are low, as one would expect, for convectively driven gravity waves typically have small phase speeds relative to the mean zonal wind where they are generated. Also the observed nondispersive behavior of the waves is easily explained by Figure 5 and points to relatively small variations of the vertical wavelength and vertical stability variations across the wave packet extending several hundred kilometers. [25] Although a full nonlinear approach is beyond the scope of this paper we note that in the limit jc – uj!0 the waves encounter a critical level where linearity does not hold and the waves break in the atmosphere. The more slowly traveling waves on Venus are therefore close to the critical level (but not breaking since we are able to observe them for 1 – 2 h in the VIRTIS frames of the same atmospheric area). Notably under the WKB approximation, as the critical level is approached the vertical wavelength diminishes and finally collapses. However, the wave stalls and the limit is only reached after a time that approaches infinite [Salby, 1996]. This process is accompanied by an increase of the wave amplitude. For the IR waves in our survey it seems that the better they can be seen the slower the phase velocities they have, what might be a signature of high-amplitude waves moving close to their critical level. [26] For the lower cloud waves we can study the double dependence of the phase speed on the two unknowns N2 (constrained from measurements from previous spacecrafts) and lz (constrained by using physically reasonable arguments). This is not feasible on the upper cloud with less observations and apparently larger variations on lz or N2. Figure 6 shows a plot of the total c2 deviation of the phase speeds of the waves from the predictions by the wave dispersion relation (equation (2a)). This was calculated for different values of N2 and lz using the equation below which also considers the estimated error in phase speeds determination. 2 N 0 2 X cx cx ðN 2 ; mÞ c N ;m ¼ ; sc i 2
ð3Þ
where the index i runs for all of the wave system measured, c0x is the measured phase speed, cx is the theoretical value from equation (2a) and sc is the variance of the measured phase speeds. [27] The contours of constant Dc2 around its minimum value can be used to determine confidence regions in the values of the varying parameters N2 and lz. A confidence region is a region within the space of parameters that contains a large percentage of the probability distribution of finding the true parameter values within it [Press at al., 1992]. If all the wave systems measured had the same values of N2 and lz, and measurement errors obeyed a
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Figure 4. Wave properties in terms of latitude. Figures 4a, 4b, 4c, and 4d show the wavelength, wave packet longitudinal extent, wave packet latitudinal extent, and the wave packet orientation relative to the parallels, respectively. Gravity waves at the upper and lower cloud level are displayed as triangles and crosses, respectively. normal error distribution, the region surrounded by the c2 = 2 curves would have a 68% probability of containing the right values of N2 and lz. The more extended regions surrounded by the c2 = 5 and c2 = 10 curves would have a 90 and 95% probability of containing the real values. From Figure 6 it seems that, when fixing N2 with the value introduced previously (N2 = 65 106 s2), the confidence region with the best fit between experiment and theory limits the range of vertical wavelengths lz between 2 and 5 km (consistent with Figure 5b). This range of vertical wavelengths matches the values of 2.5 km found from temperature scintillations measured by Magellan above the middle cloud deck [Hinson and Jenkins, 1995]. These vertically short waves have vertical wave numbers m much larger than the horizontal wave numbers k and equation (2a) can be further simplified to cx
N ; m
ð4Þ
which results in nondispersive horizontally traveling waves in accordance with the observations. [28] Convection seems the most plausible source of gravity waves in the lower atmosphere of Venus [Gierasch et al., 1997]. The possibility of internal gravity waves generated by penetrative convection has been previously considered in the terrestrial case [Stull, 1976], and it has been extensively studied in Venus through several theoretical and numerical works [Leroy and Ingersoll, 1995; Baker et al., 1998, 2000a, 2000b; Yamamoto, 2003; McGouldrick and Toon, 2008; Baker et al., 1998] which showed that compressible convection in the Venus atmosphere can lead to a significant penetration of the stable layer between 30 and 48 km both by upward convective movements from the unstable layer between 20 and 30 km, and by downward convective movements from the unstable layer between 48 and 55 km. They also pointed out the possibility that the latter convective downwellings could also affect the stable region above 55 km via convective entrainment. Baker et al. [2000b] showed that this mechanism works better
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producing enhanced overshooting of the convection in regions with high vertical wind shear, which may allow the interaction of upward and downward plumes within the 30– 48 km stable layer and generate gravity waves of higher wavelengths. In that case we could anticipate a
Figure 6. Chi-square tests exploring the space parameter defined by equation (2a) on the lower cloud waves. Confidence regions corresponding to values of chi-square larger than a given value are displayed with contour lines. The horizontal gray area indicates the range of possible buoyancy frequencies at the cloud waves’ location, and the vertical gray area indicates the corresponding range of vertical wavelengths within the confidence region of the best fit (bounded by c2 = 2 lines).
certain degree of correlation between the locations of gravity waves in the upper and lower clouds. Unfortunately the scarce gravity waves observed in the upper cloud layer may not be enough to closely examine this issue. Also, since the lower clouds are observed in the nightside hemisphere and the upper cloud are observed in the dayside hemisphere, our VIRTIS observations do not cover the same locations in high resolution and close in time.
Figure 5. Internal gravity wave dispersion relation applied to the cloud waves in the (a) upper cloud and (b) lower cloud. Phase velocities are relative to the zonal wind. Lines represent solutions to equation (2a) varying the vertical wavelengths in the ranges lz = 2 – 15 km for both clouds. We used a fixed value of the buoyancy frequency representative of the two cloud levels (N2 = 260 106 s2 for the upper cloud and N2 = 65 106 s2 for the lower cloud; details are given in section 4). 9 of 12
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Figure 7. Waves’ phase velocity and mean zonal wind profiles at the two cloud levels. Wave phase velocities and the mean zonal wind profile at the upper cloud level are displayed with dark triangles and continuous line. Wave phase velocities and the mean zonal wind profile at the lower cloud level are displayed with light crosses and discontinuous line. Those waves with no dynamical information are also displayed along a discontinuous line at zero zonal velocity. The error bars in the zonal wind profiles represent measurement errors and variance of the wind at each latitude [Sa´nchez-Lavega et al., 2008]; the error bars of the wave phases represent estimated measurement error in each wave packet.
[29] Our observations, coupled to data from previous spacecraft probes that extend in time and cover ample locations (Pioneer-Venus, Venera and Vega landers, see, e.g., Gierasch et al. [1997]), suggests that the presence of the two stable layers where the clouds lie are a persistent feature of Venus’s vertical cloud structure. Our gravity wave detections in the south hemisphere from latitudes 40°S to 75°S imply that the layer sampled with infrared images favors the vertical stability. The region equatorward of 40°S is not sampled well enough to extend this conclusion and could be dominated by stronger convective motions, whereas poleward of 75° the dipole feature may mask the presence of the waves. Although abundant and recurrent, the waves observed in the lower cloud are not permanent. In view of the stability measurements obtained from the
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different previous probes we interpret this variability as a perturbation distinct to changes in the atmospheric stability and affecting the mechanisms responsible of exciting the waves. [30] The relation between wave location and wind is examined in Figure 7. Here we show the zonal wind profiles in the upper and lower cloud layers obtained by extensive analysis of the VIRTIS data set [Sa´nchez-Lavega et al., 2008] compared to the zonal phase speeds of the wave systems in both levels. The position of the waves is correlated neither with changes in the zonal profile, nor with values of the vertical wind shear. Additionally, the width of the wave packets does not correlate with the meridional shear of the wind, what may imply that the packet width could be related with meridional variations in the atmospheric static stability. [31] The significantly larger phase speeds of the waves in the upper cloud is probably a consequence of the stronger stability in this cloud level, with much higher values for the buoyancy frequency [Seiff et al., 1985; Gierasch et al., 1997]. Contrary to the lower cloud waves, which are vertically confined in a duct of static stability, these upper cloud waves are not bounded vertically. This may result in vertical propagation to levels where they are critically absorbed by the vertically varying mean flow, or dissipated [Schubert and Walterscheid, 1984]. The absorbed waves transport energy and momentum. Their role in powering the atmospheric superrotation was examined by Leroy and Ingersoll [1995] who concluded that the waves are probably not responsible for the atmospheric global winds. [32] Most of the numerical studies of gravity waves on Venus focus on waves of smaller wavelengths (no larger than 30 km, see Baker et al. [1998, 2000a, 2000b]) than those here reported (typically 100 km). While our observations cannot distinguish these high-frequency waves (our best resolution images have a pixel size of 15– 20 km) we do detect significant wave activity of larger wavelengths (100 km). McGouldrick and Toon [2008] suggest that gravity waves with greater horizontal wavelengths comparable to our observations could be generated by broad convective plumes or a large magnitude of the vertical shear of the zonal wind but we do not observe significant convection on the upper or lower clouds. The vertical wind shear of the wind from VIRTIS data is mainly confined to the 61 to 66 km altitude range with values @ < u >/@z = 8 ± 2 ms1 per km at equatorial to midlatitudes and less than 2 ms1 per km at subpolar latitudes [Sa´nchez-Lavega et al., 2008]. [33] Gravity waves transmit energy and momentum perpendicularly to the direction of phase propagation in the direction of constant phase lines (considering the traveling direction, mainly zonal, and the vertical). Most of the waves are zonally aligned but some others are tilted and may transport some momentum meridionally. In the zonal direction some of the waves move slightly faster while others slightly more slowly than the average wind. In the meridional direction some waves move slightly northward and some other slightly southward. We also do not have information about the vertical component of the wave motion. For these reasons it was difficult to evaluate the
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average role of the waves in the global transport of energy and momentum in the atmosphere.
5. Conclusions [34] We have presented a study of the mesoscale waves present in the atmosphere of Venus at different cloud levels by using images at three wavelengths obtained by the VIRTIS instrument onboard Venus Express. We interpret them as gravity waves, and this is the first time that mesoscale waves within three different cloud levels are characterized in Venus at the same time, these levels corresponding to wavelengths 380 nm, 980 nm, and 1.74 mm. [35] The abundant waves observed in IR images display a large diversity of morphologies and wave properties with many observations revealing wave activity at a wide range of latitudes (40°S75°S). None of the wave properties seem to vary systematically with latitude or the zonal wind profile except from an observational bias of the VIRTIS data set that is unable to observe these waves at tropical and equatorial latitudes. In particular the location of the waves seems uncorrelated with topographic features on the surface or the local time that combined with the negligible effect of the solar heating at the lower cloud could indicate a crucial role of deep convection as the source of wave activity. [36] We have presented a first account on wave properties using a classical ‘‘textbook’’ linear model of gravity waves [Holton, 1992] that confirms the nature of the waves. From our statistical analysis of the wave dependence on the static stability and vertical wavelength (Figures 5b and 6) we concluded that the vertical wavelengths are on the order of 2 to 5 km in the lower cloud and 5 to 15 km in the upper cloud from the less numerous measurements of wavelength and phase speeds (Figure 5a). These results are consistent with previous determinations of the vertical wavelengths of temperature scintillations in the middle cloud measured by the Magellan spacecraft [Hinson and Jenkins, 1995]. [37] This study opens questions about the specific source of the wave activity observed, the latitudinal and temporal distributions of atmospheric stability and convection and the role of these waves in the overall dynamics of Venus atmosphere. [38] Acknowledgments. This work has been funded by Spanish MEC AYA2006-07735 with FEDER support and Grupos Gobierno Vasco IT-464-07. J.P. acknowledges a UPV fellowship, and R.H. acknowledges a ‘‘Ramo´n y Cajal’’ contract from MEC.
References Baker, R. D., G. Schubert, and P. W. Jones (1998), Cloud-level penetrative compressible convection in the Venus atmosphere, J. Atmos. Sci., 55, 3 – 18, doi:10.1175/1520-0469 (1998)055<0003:CLPCCI>2.0.CO;2. Baker, R. D., G. Schubert, and P. W. Jones (2000a), Convectively generated inertial gravity waves in the lower atmosphere of Venus. Part I: No wind shear, J. Atmos. Sci., 57, 184 – 199, doi:10.1175/15200469(2000)057<0184:CGIGWI>2.0.CO;2. Baker, R. D., G. Schubert, and P. W. Jones (2000b), Convectively generated inertial gravity waves in the lower atmosphere of Venus. Part II: Mean wind shear and wave-mean flow interaction, J. Atmos. Sci., 57, 200 – 215, doi:10.1175/1520-0469(2000)057<0200:CGIGWI>2.0.CO;2. Belton, M. J. S., G. R. Smith, G. Schubert, and A. D. Del Genio (1976), Cloud patterns, waves and convection in the Venus atmosphere, J. Atmos. Sci., 33, 1394 – 1417, doi:10.1175/1520-0469(1976)033<1394:CPWACI> 2.0.CO;2.
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Belton, M. J. S., et al. (1991), Images from Galileo of the Venus cloud deck, Science, 253, 1531 – 1536, doi:10.1126/science.253.5027.1531. Carlson, R. W., K. H. Baines, L. W. Kamp, P. R. Weissman, W. D. Smythe, A. C. Ocampo, T. V. Johnson, D. L. Matson, J. B. Pollack, and D. Grinspoon (1991), Galileo infrared imaging spectroscopy measurements at Venus, Science, 253, 1541 – 1548, doi:10.1126/science.253.5027.1541. Counselman, C. C., S. A. Gourevitch, R. W. King, G. B. Loriot, and E. S. Ginsberg (1980), Zonal and meridional circulation of the lower atmosphere of Venus determined by radio interferometry, J. Geophys. Res., 85, 8026 – 8030, doi:10.1029/JA085iA13p08026. Crisp, D., S. McMuldroch, S. K. Stephens, W. M. Sinton, B. Ragent, K. W. Hodapp, R. G. Probst, L. R. Doyle, D. A. Allen, and J. Elias (1991), Ground-based near-infrared imaging observations of Venus during the Galileo encounter, Science, 253, 1538 – 1541, doi:10.1126/ science.253.5027.1538. Del Genio, A. D., and W. B. Rossow (1990), Planetary-scale waves and the cyclic nature of the cloud top dynamics on Venus, J. Atmos. Sci., 47, 293 – 318, doi:10.1175/1520-0469(1990)047<0293:PSWATC>2.0.CO;2. Drossart, P., et al. (2007), Scientific goals for the observation of Venus by VIRTIS on ESA/Venus express mission, Planet. Space Sci., 55, 1653 – 1672, doi:10.1016/j.pss.2007.01.003. Esposito, L. W., J. L. Bertaux, V. Krasnopolsky, V. I. Moroz, and L. V. Zasova (1997), Chemistry of lower atmosphere and clouds, in VENUS II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, pp. 415 – 458, Univ. of Ariz. Press, Tucson. Flasar, F. M., and P. J. Gierasch (1986), Mesoscale waves as a probe of Jupiter’s deep atmosphere, J. Atmos. Sci., 43, 2683 – 2707, doi:10.1175/ 1520-0469(1986)043<2683:MWAAPO>2.0.CO;2. Gierasch, P. J. (1987), Waves in the atmosphere of Venus, Nature, 328, 510 – 512, doi:10.1038/328510a0. Gierasch, P. J., et al. (1997), The general circulation of the Venus atmosphere: An assessment, in VENUS II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, pp. 459 – 500, Univ. of Ariz. Press, Tucson. Hinson, D. P., and J. M. Jenkins (1995), Magellan radio occultation measurements of atmospheric waves on Venus, Icarus, 114(2), 310 – 327, doi:10.1006/icar.1995.1064. Holton, J. R. (1992), An Introduction to Dynamic Meteorology, 535 pp., Elsevier, New York. Hou, A. Y., and B. F. Farrell (1987), Superrotation induced by critical-level absorption of gravity waves on Venus: An assessment, J. Atmos. Sci., 44, 1049 – 1061, doi:10.1175/1520-0469(1987)044<1049:SIBCLA> 2.0.CO;2. Houze, R. A. (1993), Cloud Dynamics, 573 pp., Elsevier, New York. Hunt, G. E., and J. P. Muller (1979), Voyager observations of small-scale waves in the equatorial region of the jovian atmosphere, Nature, 280, 778 – 780, doi:10.1038/280778a0. Kasprzak, W. T., A. E. Hedin, H. G. Mayr, and H. B. Niemann (1988), Wavelike perturbations observed in the neutral thermosphere of Venus, J. Geophys. Res., 93, 11,237 – 11,245, doi:10.1029/JA093iA10p11237. Kliore, A. J., and I. R. Patel (1980), Vertical structure of the atmosphere of Venus from Pioneer Venus orbiter radio occultations, J. Geophys. Res., 85, 7957 – 7962, doi:10.1029/JA085iA13p07957. Kolosov, M. A., O. I. Yakovlev, A. I. Efimov, S. S. Matyugov, T. S. Timofeeva, E. V. Chub, A. G. Pavelyev, A. I. Kucheryavenkov, I. E. Kalashnikov, and O. E. Milekhin (1980), Investigation of the Venus atmosphere and surface by the method of radiosounding using VENERA-9 and 10 satellites, Acta Astronaut., 7, 219 – 234, doi:10.1016/ 0094-5765(80)90062-4. Leroy, S. S., and A. P. Ingersoll (1995), Convective generation of gravity waves in Venus’s atmosphere: Gravity wave spectrum and momentum transport, J. Atmos. Sci., 52, 3717 – 3737, doi:10.1175/1520-0469(1995) 052<3717:CGOGWI>2.0.CO;2. Leroy, S. S., and A. P. Ingersoll (1996), Radio scintillations in Venus’s atmosphere: Application of a theory of gravity wave generation, J. Atmos. Sci., 53, 1018 – 1028, doi:10.1175/1520-0469(1996)053<1018:RSIVAA> 2.0.CO;2. Markiewicz, W. J., D. V. Titov, S. S. Limaye, H. U. Keller, N. Ignatiev, R. Jaumann, N. Thomas, H. Michalik, R. Moissl, and P. Russo (2007), Morphology and dynamics of the upper cloud layer of Venus, Nature, 450, 633 – 636, doi:10.1038/nature06320. McGouldrick, K., and O. B. Toon (2008), Observable effects of convection and gravity waves on the Venus condensational cloud, Planet. Space Sci., 56, 1112 – 1131, doi:10.1016/j.pss.2008.02.010. Niemann, H. B., W. T. Kasprzak, A. E. Hedin, D. M. Hunten, and N. W. Spencer (1980), Mass spectrometric measurements of the neutral gas composition of the thermosphere and exosphere of Venus, J. Geophys. Res., 85, 7817 – 7827, doi:10.1029/JA085iA13p07817. Peralta, J., R. Hueso, and A. Sa´nchez-Lavega (2007a), Cloud brightness distribution and turbulence in Venus using Galileo violet images, Icarus, 188, 305 – 314, doi:10.1016/j.icarus.2006.12.005.
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Peralta, J., R. Hueso, and A. Sa´nchez-Lavega (2007b), A reanalysis of Venus winds at two cloud levels from Galileo SSI images, Icarus, 190, 469 – 477, doi:10.1016/j.icarus.2007.03.028. Piccioni, G., et al. (2007), South-polar features on Venus similar to those near the north pole, Nature, 450, 637 – 640, doi:10.1038/nature06209. Press, W. H., S. A. Teukolsky, W. T. Vetterling, and B. P. Flannery (1992), Numerical Recipes in Fortran 77: The Art of Scientific Computing, pp. 687 – 693, Cambridge Univ. Press, New York. Reuter, D. C., et al. (2007), Jupiter cloud composition, stratification, convection and wave motion: A view from New Horizons, Science, 318, 223 – 225, doi:10.1126/science.1147618. Rossow, W. B., et al. (1980), Cloud morphology and motions from Pioneer Venus images, J. Geophys. Res., 85, 8107 – 8128, doi:10.1029/ JA085iA13p08107. Sagdeev, R. Z., V. M. Linkin, J. E. Blamont, and R. A. Preston (1986), The VEGA Venus balloon experiment, Nature, 328, 510 – 512. Salby, M. L. (1996), Fundamentals of Atmospheric Physics, 627 pp., Elsevier, New York. Sa´nchez-Lavega, A., et al. (2008), Variable winds on Venus mapped in three dimensions, Geophys. Res. Lett., 35, L13204, doi:10.1029/ 2008GL033817. Saunders, R. S., R. E. Arvidson, J. W. Head III, G. G. Schaber, E. R. Stofan, and S. C. Solomon (1991), An overview of Venus geology, Science, 252, 249 – 252, doi:10.1126/science.252.5003.249. Schubert, G. (1983), General circulation and the dynamical state of the Venus, in VENUS, pp. 681 – 765, Univ. of Ariz. Press, Tucson. Schubert, G., and R. L. Walterscheid (1984), Propagation of small-scale acoustic-gravity waves in the Venus atmosphere, J. Atmos. Sci., 41, 1202 – 1213. Seiff, A., D. B. Kirk, R. E. Young, R. C. Blanchard, J. T. Findlay, G. M. Kelly, and S. C. Sommer (1980), Measurements of thermal structure and thermal contrasts in the atmosphere of Venus and related dynamical observations: Results from the four Pioneer Venus probes, J. Geophys. Res., 85, 7903 – 7933, doi:10.1029/JA085iA13p07903. Seiff, A., J. T. Schofield, A. J. Kliore, F. W. Taylor, S. S. Limaye, H. E. Revercomb, L. A. Sromovsky, V. V. Kerzhanovich, V. I. Moroz, and M. Y. Marov (1985), Models of the structure of the atmosphere of
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Venus from the surface to 100 kilometers altitude, Adv. Space Res., 5(11), 3 – 58, doi:10.1016/0273-1177(85)90197-8. Stull, R. B. (1976), Internal gravity waves generated by penetrative convection, J. Atmos. Sci., 33, 1279 – 1286, doi:10.1175/15200469(1976)033<1279:IGWGBP>2.0.CO;2. Svedhem, H., et al. (2007), Venus Express: The first European mission to Venus, Planet. Space Sci., 55, 1636 – 1652, doi:10.1016/j.pss.2007. 01.013. Taylor, F. W., et al. (1980), Structure and meteorology of the middle atmosphere of Venus: Infrared remote sensing From the Pioneer Orbiter, J. Geophys. Res., 85, 7963 – 8006, doi:10.1029/JA085iA13p07963. Toigo, A., P. J. Gierasch, and M. D. Smith (1994), High resolution cloud feature tracking on Venus by Galileo, Icarus, 109, 318 – 336, doi:10.1006/icar.1994.1097. Yamamoto, M. (2003), Gravity waves and convection cells resulting from feedback heating of Venus’ lower clouds, J. Meteorol. Soc. Jpn., 81, 885 – 892, doi:10.2151/jmsj.81.885. Young, R. E., R. L. Walterscheid, G. Schubert, A. Seiff, V. M. Linkin, and A. N. Lipatov (1987), Characteristics of gravity waves generated by surface topography on Venus: Comparison with the VEGA balloon results, J. Atmos. Sci., 44, 2628 – 2639, doi:10.1175/15200469(1987)044<2628:COGWGB>2.0.CO;2. Young, R. E., R. L. Walterscheid, G. Schubert, L. Pfister, H. Houben, and D. L. Bindschadler (1994), Characteristics of finite amplitude stationary gravity waves in the atmosphere of Venus, J. Atmos. Sci., 51, 1857 – 1875, doi:10.1175/1520-0469(1994)051<1857:COFASG>2.0.CO;2.
P. Drossart, Observatoire de Paris, LESIA, 5 place Jules Janssen, F-92195 Meudon, France. R. Hueso, J. Peralta, and A. Sa´nchez-Lavega, Departamento de Fı´sica Aplicada I, Escuela Te´cnica Superior de Ingenierı´os, Universidad del Paı´s Vasco, Alameda Urquijo s/n, E-48013, Bilbao, Spain. (javier.peralta@ ehu.es) G. Piccioni and O. Lanciano, Istituto di Astrofisica Spaziale e Fisica Cosmica, INAF, via del Fosso del Cavaliere 100, I-00133 Rome, Italy.
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Venus Express/VIRTIS observations of middle and lower cloud variability and implications for dynamics K. McGouldrick,1 K. H. Baines,2 T. W. Momary,2 and D. H. Grinspoon1 Received 1 February 2008; revised 27 June 2008; accepted 20 August 2008; published 18 November 2008.
[1] We present an analysis of Venus Express Visible and Infrared Thermal Imaging
Spectrometer (VIRTIS) data, carried out to characterize the morphological, geographical, and evolutionary trends of the middle and lower cloud features that are observed in the atmosphere of Venus as variations in brightness temperatures in specific near-infrared wavelengths. In this preliminary study, we analyze only data collected over the span of 11 orbits. The mean radiance as a function of latitude is consistent with previous ground-based observations, indicating that the overall global distribution of mean cloud cover is stable, at least on a 10- to 20-year time scale. In contrast with the consistent level of radiance at high latitudes, a significant amount of variability to the radiance exists at lower latitudes, consistent with significant convective activity in the lower and middle cloud decks. The morphology of the holes tends from highly variable orientations of features with aspect ratios of nearly one at low latitudes, to very large aspect ratios and zonally oriented features at higher latitudes. The peak radiance of the holes appears not to demonstrate a latitudinal tendency. There is evidence of more variability to the morphology and radiance of features at lower latitudes. To investigate the evolution of the holes, we examine a sequence of images taken over a 5 h span of a single orbit. If this limited amount of data is representative, then the typical e-folding time scale for the evolution of a hole is about 1 day. Citation: McGouldrick, K., K. H. Baines, T. W. Momary, and D. H. Grinspoon (2008), Venus Express/VIRTIS observations of middle and lower cloud variability and implications for dynamics, J. Geophys. Res., 113, E00B14, doi:10.1029/2008JE003113.
1. Introduction [2] Holes in the middle and lower cloud decks of Venus were first discovered as spatial inhomogeneities in the nearinfrared brightness temperatures [Allen and Crawford, 1984]. Thermal emission from the lower atmosphere and surface of Venus escapes through narrow spectral windows among the near-infrared absorption bands of carbon dioxide and water vapor. This emission is scattered by the sulfuric acid cloud decks. From tracking of the features as they traversed the planet and detailed radiative transfer analysis of the spectra of those features, it was recognized that they represented variations in the lower and middle cloud opacity [Crisp et al., 1989]. [3] The clouds are a defining characteristic of the atmosphere of Venus. They completely enshroud the planet between the altitudes of about 50 and 70 km. Consequently, except for the variations noted in the previous paragraph, the clouds render the lower atmosphere difficult to observe in infrared, visible, and ultraviolet wavelengths. The clouds of Venus have been implicated as playing a significant role 1 Department of Space Sciences, Denver Museum of Nature and Science, Denver, Colorado, USA. 2 Jet Propulsion Laboratory, California Institute of Technology, Pasadena, California, USA.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003113$09.00
in the maintenance of the Venus greenhouse effect [Pollack et al., 1980]. Thus, breaks in the clouds, which allow a greater flux of radiation to escape (and absorb a smaller amount of solar radiation), can have a significant effect on the thermal balance of Venus. This can have significant ramifications, especially in terms of the long-term climate evolution of Venus [Crisp, 1989; Hashimoto and Abe, 2000; Bullock and Grinspoon, 2001]. [4] The atmosphere of Venus exhibits a global superrotation at most altitudes. In order to sustain the global superrotation, there must be vertical transport of horizontal momentum to the upper atmosphere, specifically between about 40 and 80 km. Vertical transport of momentum by eddies and by waves has been suggested [Schubert et al., 1980]. However, both the eddies and the waves are going to interact in some way with the clouds that occupy the altitudes between 50 and 70 km. Most likely, vertically traveling waves and eddies will be dissipated by the regions of instability in the vicinity of the clouds. The radiative dynamical feedback that supports the cloud [Pollack et al., 1980; McGouldrick and Toon, 2007], maintains a region of dynamic instability in the middle and lower cloud region, which can limit the efficiency of vertical momentum transport by wave propagation. The shear profile and the static stability profile, both of which affect the vertical transport of energy that is brought about by convection, will have the largest effect on what happens to these vertically propagating waves and eddies. Furthermore, the morphological appear-
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ance of cloud features (aspect ratio and ‘‘tilt’’ relative to the zonal direction) can offer information on the nature of these atmospheric characteristics (shear profile, static stability profile, and the amount of convective overturning). Thus, the morphology of the clouds can provide insight to the dynamical environment; hence, the efficiency of vertical momentum transport by these processes. [5] For these reasons, an understanding of what sustains the clouds, and what sorts of changes are responsible for the formation of holes in those clouds, is necessary in order to understand the Venus atmosphere as a whole. In this paper, we compare the results and predictions of several recent theoretical studies of the clouds of Venus, to recent observations made by the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) on Venus Express. We analyze the possible existence of latitudinal tendencies of cloud morphology and overall cloud opacity, as well as the typical cloud feature lifetimes. 1.1. Previous Observations of the Clouds and Holes [6] Crisp et al. [1989] showed that the near-infrared features discovered by Allen and Crawford [1984] were caused by variations in the opacity of the middle and lower cloud decks. The observations by Crisp et al. [1989] also provided an early general characterization of the morphology and distribution of the cloud features at wavelengths of 1.74 mm and 2.3 mm in the near infrared. Crisp et al. [1989] observed an apparent latitudinal dependence of the overall cloud opacity, in which the opacity was greatest at high latitudes, and a relatively featureless bright band associated with diminished cloud opacity that often existed at mid latitudes (around 40° to 60°). Crisp et al. [1989] also noted occasional significant asymmetry between the brightness (hence, opacity) of the northern and southern hemispheres. [7] The flyby of the Galileo spacecraft in February of 1990 provided an excellent opportunity for the study of these near-infrared cloud features. Belton et al. [1991] observed the clouds of Venus with the Galileo solid-state imaging (SSI) camera. It is important to note, however, that the features observed by SSI likely are due to cloud variability at a higher altitude (i.e., about 60 to 65 km) than the cloud variability observed in the near-infrared spectral windows on the nightside (i.e., around 50 to 55 km). This is partly because the dayside observations detected solar infrared radiation that has been reflected from near the tops of the clouds, whereas the nightside observations detected surface-emitted infrared radiation transmitted through the clouds from below. Gaseous absorption by the atmosphere extinguishes the emitted or reflected radiation at much higher altitudes outside of the narrow spectral windows. Thus, the broadband observation exhibits contributions from a greater range of altitudes (mostly higher) than does the narrow-band observation. Nevertheless, Belton et al. [1991] noted that cloud features tended to appear ‘‘patchy’’ at low latitudes and ‘‘streaky’’ at high latitudes. Viewing the cloudy atmosphere at spatial resolutions as fine as 17 km per pixel, Belton et al. [1991] detected changes in cloud morphology on time scales only in excess of an hour or two; generally on the order of 1 day. They also noted the presence of ‘‘bright-rimmed cellular features,’’ suggestive of cellular convective elements that had lifetimes in excess of a day, and sizes on the order of a few hundred kilometers.
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Blamont et al. [1986] also noted a tendency of cloud features at higher latitudes to tilt relative to the zonal direction in a way that suggested a spiral pattern. Such a spiral (i.e., the polar vortex) is known to exist over both the north and south poles of Venus [Taylor et al., 1979; Piccioni et al., 2007]. The latitudinal extent of the Piccioni et al. [2007] observations is only poleward of about 50° latitude. Analysis of the tilting of cloud features in images centered over middle and lower latitudes can place constraints on the equatorward influence of these polar vortices. [8] Carlson et al. [1991], observing the nightside of Venus with the Galileo Near-Infrared Imaging Spectrometer (NIMS), noted that contrast between the brightest and darkest regions of the nightside of Venus at 1.74 mm was about 5:1. Carlson et al. [1991] also noted the existence of linear features whose tilt relative to the parallels of latitude implied meridional transport, presumably by Hadley circulation. These features seem to be similar in appearance to those described by Blamont et al. [1986] and Belton et al. [1991]; but Carlson et al. [1991] noted that establishment of a direct correlation of features between the two instruments (NIMS and SSI) was uncertain. [9] While the Galileo Near-Infrared Mapping Spectrometer (NIMS) was able to obtain only, essentially, a snapshot of the planet during the flyby, ground-based observations by Crisp et al. [1991a], spanning the weeks leading up to and following the flyby, were able to provide some context for the NIMS observations. Crisp et al. [1991a] detected a latitudinal distribution of cloud opacity similar to that observed by Crisp et al. [1989]. That is, they noted that high-latitude regions tended to be ‘‘dark and featureless,’’ whereas midlatitude regions tended to be ‘‘occupied by bright quasi-zonal bands.’’ In addition, low latitudes were noted to exhibit the most variability, both in terms of feature opacity and size. Crisp et al. [1991a] also noted the existence of a very large (hemisphere-scale) dark region that had persisted throughout the duration of the observing sessions (about 7 weeks). Chanover et al. [1998] also noted the existence of such a hemispherical asymmetry. Finally, Crisp et al. [1991a] noted that the features that were visible to Galileo during the flyby were of a relatively small scale that tended not to persist for more than about a week (i.e., about one atmospheric rotational period). [10] Markiewicz et al. [2007b] described observations of Venus cloud structure observed with the Venus Monitoring Camera (VMC) [Markiewicz et al., 2007a] on board Venus Express. While the infrared channels of VMC can be used to probe the middle and lower clouds of Venus, the observations reported by Markiewicz et al. [2007b] are mostly in the ultraviolet channel. Thus, their observations are of cloud features located largely above 65 km altitude. They observe the polar vortex as a dark spiral structure in the ultraviolet images. They observe a mottled cloud structure at low latitudes, which they attribute to vigorous convection in the upper cloud deck, with the sizes of typical convective cells being about 30 km. They also report waves with wavelengths on the order of 10 to 30 km in the upper cloud region. 1.2. Previous Simulations of the Clouds and Holes [11] Simulations suggest that the global dynamics of Venus have a significant effect on the formation and distribution of the Venus clouds and their holes. Imamura
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and Hashimoto [1998] showed that the Hadley circulation could result in enhancements of cloud opacity in the vicinity of both the ascending branch and the descending branch of the near-global Hadley cell. McGouldrick and Toon [2008b] showed that differences in the structure of the vertical shear of the zonal winds could have an effect on the evolution of large (2000 km) holes in the clouds of Venus. McGouldrick and Toon [2008a] showed that convective cells with typical size and separation of about 30 km tended to generate holes with lifetimes on the order of an hour or two, and with optical depths that were consistent with observations. McGouldrick and Toon [2008a] also showed that gravity waves, with characteristics similar to those of waves putatively launched by convection that have been both observed [Hinson and Jenkins, 1995] and simulated [Baker et al., 2000], can generate wave-like features in the clouds with 10 to 30 km wavelengths, visible in the emitted 1.74 mm radiation. [12] Imamura and Hashimoto [1998] suggested that the Hadley circulation determines the latitudinal variation in the middle and lower cloud deck opacity. Upwelling in the ascending branch near the equator leads to enhanced opacity because of the influx of sulfuric acid vapor into the cloud region. Meridional transport toward the descending branch of the Hadley circulation, which they reasoned to occur poleward of 55° latitude, also leads to enhanced cloud opacity because of an accumulation of upper cloud mass that has been transported poleward by the upper level Hadley circulation (because the time scale for the sedimentation of the upper cloud particles was considerably longer than the time scale for poleward transport of those particles by the Hadley circulation). Furthermore, the bright region between about 40° and 60° latitude could be explained as an absence of these cloud-thickening mechanisms. They also found that if the Hadley circulation were slower, then the contrast between the brightness at midlatitudes and that at high latitudes would be less severe, as a result of a thinning of the upper cloud at high latitudes. If the Hadley circulation were less efficient, then the time for the upper level flow to travel from equator to pole no longer would be large compared with sedimentation times of the upper cloud particles. Hence, less photochemically produced mass is transported to high latitudes. They also found that the bright band was located more equatorward in the case of a less efficient Hadley circulation. Hence, it may be possible to determine the magnitude of the Hadley circulation by determining the latitudinal location of the high-latitude peak in total cloud opacity [Imamura and Hashimoto, 1998]. [13] The middle and lower cloud decks of Venus are sustained by a radiative dynamical feedback, whereby heating of the cloud base by the very warm deep atmosphere drives an unstable lapse rate within the region occupied by the middle and lower cloud decks [McGouldrick and Toon, 2007]. Furthermore, the time scales for growth and for dynamical transport (by winds, and by convection, when present) are much shorter than the radiative time scale. Hence, in the simulations by McGouldrick and Toon [2007], the structure of the clouds was dominated by brief, intense periods of vertical transport (i.e., convection), separated by longer periods of quiescence during which the thermal instability developed.
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[14] McGouldrick and Toon [2008b] suggested that the nature of the vertical wind profile affects the longevity and possibly the morphology of large holes (2000 km) in the clouds. Specifically, the vertical shear of the zonal wind, coupled with the radiative dynamical feedback, can effectively dissipate large holes in the clouds. Furthermore, the nature of the vertical shear profile affected the time required to smooth out a simulated hole in the clouds. In simulations exhibiting negligible vertical shear throughout the middle cloud deck (as have been observed at low latitudes and midlatitudes by Pioneer Venus [Counselman et al., 1980] and measured via multispectral wind tracking with VIRTIS on board Venus Express [Sa´nchez-Lavega et al., 2008]), the lifetimes of holes were longer than in simulations in which the shear was at a constant magnitude throughout the entire cloud domain (i.e., shear similar to what has been observed at higher latitudes, especially by the Pioneer Venus north probe). Although the radiative dynamical feedback was able to limit the zonal ‘‘stretching’’ of holes in the simulations by McGouldrick and Toon [2008b] that exhibited minimal shear within the middle and lower cloud decks, the holes were stretched out across large zonal distances in simulations that exhibited greater amounts of shear. Such differences in the aspect ratios of features in the Venus clouds observed by VIRTIS could indicate the vertical structure of the zonal winds in the region of the middle and lower cloud decks of Venus. That is, holes that become highly zonally elongated may be in regions of high vertical shear of the zonal wind; and holes that maintain a more blocky appearance may be in regions of low vertical shear of the zonal wind. [15] An alternative cause for this zonal stretching of cloud features could be horizontal shear of the zonal wind. However, observations by both the Venus Monitoring Camera and VIRTIS on Venus Express indicate that there is negligible meridional shear of the zonal wind equatorward of about 50° latitude [Markiewicz et al., 2007b; Sa´nchez-Lavega et al., 2008]. Thus, at latitudes equatorward of 50°, the only contribution to stretching by the meridional shear of the zonal wind arises from the change in planetary circumference with latitude (f): @w 86400 u tan f ¼ ; @f RV cos f
ð1Þ
where RV = 6100 103 m is the radius of the planet at the altitude of the clouds, w is the angular velocity of the atmosphere in units of ° longitude per day, and the factor 86400 is needed to convert the units of the RHS to day1. For f 50°, and u = 60 m s1 from Sa´nchez-Lavega et al. [2008], this rate does not exceed 1.6° longitude per day per degree of latitude. Thus, among features at midlatitudes to low latitudes, only those holes that persist for several days, can be stretched significantly by the meridional shear. Furthermore, the stretching of the cloud features cannot be the result of zonal shear of the zonal wind. Such a situation would result in horizontal convergence or divergence, which would force vertical motion of atmospheric parcels. Such vertical motion can create significant changes in cloud opacity [McGouldrick and Toon, 2007, 2008a], but will not cause a zonal elongation of the features. [16] McGouldrick and Toon [2008a] analyzed the effect of gravity waves and convection cells on the appearance of
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Table 1. Data Analyzed in Morphological Analyses File Name
Orbit
VI0363_00 VI0364_00 VI0365_00 VI0366_00 VI0367_00 VI0368_00 VI0369_00 VI0370_00 VI0371_00 VI0372_00
363 364 365 366 367 368 369 370 371 372
Start Date 18 19 20 21 22 23 24 25 26 27
Apr Apr Apr Apr Apr Apr Apr Apr Apr Apr
2007 2007 2007 2007 2007 2007 2007 2007 2007 2007
Start Timea
Exposure Timeb (s)
Distance (km)
2017:56 2017:10 2016:25 2015:42 2014:57 2014:12 2013:25 2012:41 2011:55 2011:10
0.36 3.3 0.36 3.3 0.36 3.3 0.36 3.3 0.36 3.3
67021.2 66775.9 66468.1 66226.3 67209.0 67226.6 67218.5 67223.6 67203.3 67220.6
a
Times and dates are UTC. b Exposure time is per slice.
the clouds of Venus. They found, from a simple kinematical model used to drive the microphysical model utilized by McGouldrick and Toon [2007], that convective cells with sizes on the order of 30 km and vertical velocities of ±2 m s1 produce optical depth variations comparable in magnitude to those that have been observed in the clouds of Venus [Grinspoon et al., 1993], and which experience lifetimes on the order of hours. These simulated convective cells are consistent with those that have been observed with the Vega balloons [Linkin et al., 1986], and that have been generated in convective simulations by Baker et al. [2000]. McGouldrick and Toon [2008a] also found that gravity waves, such as those found in the simulations of Baker et al. [2000], and in the occultation observations by Hinson and Jenkins [1995], would lead to the appearance of wave-like structures in the lower clouds of Venus, with wavelengths comparable to the size of the convective cells. 1.3. VIRTIS [17] The VIRTIS-M IR channel is a medium (spatial) resolution (0.25 mrad) infrared mapping spectrometer that covers the range of wavelengths from 1 mm to 5 mm, with a spectral resolution of roughly l/Dl250. This is sufficient spectral resolution to resolve the spectral windows [Drossart et al., 2007]; and the spatial resolution allows for the detection of features as small as about 30 km when the spacecraft is at apoapse (d66,000 km). When the spacecraft is within about 10,000 km of the planet during the course of its highly eccentric orbit, it is moving too quickly for VIRTIS to operate very effectively in its mapping modes [Drossart et al., 2007]. Thus, all of the data we analyze in this paper are obtained at spacecraft distances greater than 10,000 km. [18] With clever pointing over the course of multiple orbits, a significant latitudinal and longitudinal coverage of the clouds of the southern hemisphere may be built up, which allows for the determination of possible latitudinal tendencies in the location of and optical thickness of the clouds. In the course of a single orbit, it is possible to observe several hours worth of cloud evolution (up to about 6 h) of a small region of the Venus atmosphere (defined by the field of view common to all of the images in the series: the maximum possible region being the total field of view of the image taken when the spacecraft is closest to the planet). Analysis of longer-term evolution is possible, but requires a repointing of the spacecraft on successive orbits, since the
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previously observed region of the atmosphere will have rotated out of the relatively narrow field of view by the time the spacecraft next arrives in position for observation by the mapping modes of VIRTIS. Even with successive repointings that follow the rotation of the atmosphere, significant changes in cloud morphology during the time that VIRTIS-M is unable to observe the planet frequently make difficult the recovery of features common to observations on successive orbits.
2. Observational Goals and Methods [19] Specific latitudinal variations in overall cloud opacity have been predicted to be a consequence of the Hadley circulation [Imamura and Hashimoto, 1998], and have been recorded by ground-based observations [Crisp et al., 1991a]. Does a large-scale latitudinal dependence of cloud cover exist in the VIRTIS data, as suggested by previous ground-based observations? Are there significant spatial or temporal variations in the latitudinal distribution of cloud cover? Tendencies in cloud shape (aspect ratio) and orientation (angle of ‘‘long axis,’’ relative to the zonal direction) can indicate differences in the shear structure or in the level of convection that is occurring. Does a latitudinal tendency to cloud shape or orientation exist in the VIRTIS data? Since the clouds represent a significant contributor to the greenhouse effect on Venus, their variability can alter the thermal balance of the planet. How variable are the clouds, and how does cloud variability affect the thermal balance? We analyze a subset of the VIRTIS-M IR observations to attempt to answer these questions concerning the interactions between the atmospheric dynamics and the cloud structure of Venus. [20] In all cases, we perform our analyses using the 1.74 mm image of the VIRTIS-M IR data. We choose this wavelength band over those in the other spectral windows because the 1.74 mm band possesses the best combination of high peak radiance and sufficient contrast between cloudy and clear observations; thus, the signal-to-noise ratio of both the brighter holes and the darker clouds is greater than it is for any other spectral image. Furthermore, while the images in the 2.3 mm window region demonstrate a greater overall brightness contrast (around 20:1 in the Galileo NIMS data, compared with a roughly 5:1 ratio at 1.74 mm [Carlson et al., 1991]), the radiance in the cloudiest regions, is barely distinguishable (and possibly indistinguishable) from the background noise. In order to analyze the latitudinal distribution of cloud cover and cloud morphology on Venus, we use images/cubes taken near the apoapse (66,000 km) of several successive orbits (orbits 363 through 372), as summarized in Table 1. We use images obtained near apoapse in order to maximize our spatial coverage of the planet, with only a small loss in spatial resolution. The data are calibrated using the VIRTIS PDS/IDL software library, as described by Drossart et al. [2007]. [21] Figure 1a demonstrates the overall coverage of the middle atmosphere by these selected observations. In generating this plot, we assume that the atmosphere rotates at a solid body rate of once per 5.5 days (a speed of about 80 m s1 at the equator, or about 60 m s1 at 40° latitude), and shift the geographic longitudes of each observation to the east, according to the time separation of each image from
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[22] Figure 1b shows the frequency of observations by location in the atmosphere. We see that among these images there is significant overlap, such that about 35% of the atmosphere between +10° latitude and 90° latitude has been imaged at least twice; and some regions (such as the one centered near 50° longitude and 40° latitude) have been observed three times or more. Such repeated coverage of the same part of the atmosphere gives hope to the possibility of ultimately extending the analysis of the variation in feature morphology and evolution from time scales no greater than 6 h (the extent of observation possible during a single orbit) to time scales on the order of days.
Figure 1. Distribution and frequency of coverage of the atmosphere by the images analyzed in this paper. (a) The regions of the atmosphere that have been imaged at least once over the course of the considered images (crosshatched regions). (b) The frequency of coverage of the atmosphere. The numbered contours indicate the number of images among those considered here that view the same region of the atmosphere. Both plots assume the atmosphere rotates as a solid body with a period of 5.5 days.
the first image (orbit 363). We choose a planet-circling period of 5.5 days for the clouds because that period is consistent with previous observations of lower and middle cloud tracking [Allen and Crawford, 1984; Crisp et al., 1991b]. According to the recent analysis of wind speeds from VIRTIS data by Sa´nchez-Lavega et al. [2008], a zonal wind speed of about 80 m s1 corresponds to an altitude between 58 and 70 km, for latitudes equatorward of 50°. A 5.5 day period at 40° latitude is consistent with the 60 m s1 winds reported by Sa´nchez-Lavega et al. [2008] at altitudes between 46 and 61 km. Though it is an oversimplification to assume solid body rotation, doing so allows us to obtain at least a first-order estimate of the atmospheric coverage of the observations. It also helps to facilitate the identification of features from image to image, despite their sometimes significant evolution (see section 2). Figure 1a shows that the images that we consider for this particular analysis provide us with coverage of a significant fraction of the southern hemisphere atmosphere (about 72% coverage of the area between the latitudes of +10° and 90°, with most of the missing regions lying poleward of 70°). This extensive coverage of the atmosphere allows us to have confidence in the conclusions that we draw regarding the latitudinal distributions of the features. If these images did not cover so much of the planet longitudinally, our analyses might have been blind to hemispherical variations such as those that have been reported by Crisp et al. [1989, 1991b].
2.1. Latitudinal Distribution of Cloud Cover and Hole Morphology [23] Crisp et al. [1989, 1991b] noted the appearance of a significant latitudinal distribution of cloud cover in which midlatitudes tended to be brighter, and higher latitudes tended to be darker. Imamura and Hashimoto [1998] provided a theoretical explanation of such a latitudinal distribution of cloud opacity, which, they determined, is caused by the action of the Hadley circulation. Here again, the tremendous volume of the VIRTIS data can play a significant role in confirming or refuting the existence and consistency of such a distribution of cloud opacity. Additionally, the VIRTIS data can be used to more finely tune such simulations of the Hadley circulation on Venus. To analyze the latitudinal distribution of cloud cover, we remap the VIRTIS data to a cylindrical projection and correct for the emission angle by applying the correction used by Carlson et al. [1993]. We then find the mean and standard deviation of the radiance as a function of latitude for each image. We also calculate a longer-term average latitudinal distribution of radiance, in which we average over all 10 images considered here. [24] Belton et al. [1991] suggested that there was a trend in the shapes of the features as a function of latitude. Specifically, the features at lower latitudes appeared ‘‘blocky,’’ whereas those at higher latitudes appeared streaky. The great volume of VIRTIS data from Venus Express can indicate whether these trends noted by Belton et al. [1991] are real and/or persistent. We investigate these possible morphological tendencies of the clouds by identifying features in each image, and determining a longitudinal versus transverse aspect ratio for each feature. The features are identified by using the latitudinal profile of the mean radiance of each image, as described in section 2.1. We subdivide the image into latitude bands that are no greater than 15° in extent. We find the mean and the standard deviation of the radiance over those latitudes in that image, and draw contours of the mean and the mean plus one standard deviation onto the image (Figure 2). A feature is identified if the mean plus standard deviation contour is closed (or nearly so) and convex. At higher latitudes, where the longitudinal variation is smaller and the latitudinal variation is greater, we use narrower bands to determine the defining contours. We determine the longitudinal versus transverse aspect ratio by measuring the distances along the long and short axes of each feature. The feature is assigned latitude and longitude coordinates that identify its left (L), right (R), top (T), and bottom (B) extent. The aspect ratio is then the ratio of the LR arc to the TB arc. Thus, a large aspect ratio indicates a highly elongated feature,
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Figure 2. (right) 1.74 mm VIRTIS image VI0383_00. (left) The blowup highlights and zooms in on the boxed region of the VIRTIS image. The parameters measured in the course of the morphological analyses are shown in the zoomed image at left. whereas an aspect ratio closer to one indicates a more blocky feature. [25] We also accumulate information on the tilt of the features relative to the parallels of latitude. The tilt angle is taken to be the angle that the LR arc of the feature makes relative to the parallels of latitude. Depending upon the magnitude of the meridional and zonal winds, a tilt measured from the parallels of latitude can indicate meridional flow or shear of cloud material or energy or a significant vertical shear of the zonal wind. 2.2. Cloud/Hole Lifetimes [26] The duration of holes in the clouds of Venus can indicate the nature of the processes that sustain them. McGouldrick and Toon [2008b] suggested that differences in the vertical profile of the zonal wind could lead to differences in the longevity of the features. Furthermore, McGouldrick and Toon [2008b] noted that a combination of vertical shear of the zonal wind and the vertical motions driven by the radiative dynamical feedback was sufficient to dissipate a large (2000 km) hole in a time scale of under two weeks. Crisp et al. [1991b] reported that a 2000 km hole persisted for at least two weeks. However, since the initial formation of that particular feature was not observed, this is only a lower limit. Furthermore, the spatial resolution of VIRTIS on Venus Express, even at apoapse, is far
superior to what Crisp et al. [1991b] were able to achieve from Earth. Thus, the analysis of the evolution of smaller features is now possible. Since the absorption of upwelling infrared radiation plays a role in sustaining the middle and lower clouds [Pollack et al., 1980; McGouldrick and Toon, 2007], the longevity of holes affects the magnitude of possible horizontal temperature or pressure variations that would result from such heating. Strong horizontal temperature or pressure gradients will affect the magnitude of vertical motions triggered by the radiative dynamical feedback. [27] To determine the typical lifetime of the cloud features, we analyze the VIRTIS-M data from orbit 383. This orbit contains a sequence of cubes that spans 5 uninterrupted h, and each cube in the sequence exhibits exposure times long enough that the nightside (rather than the dayside) is exposed ideally (Table 2). Since the Venus Express spacecraft is approaching Venus throughout the series of acquired cubes in each orbit, there is only a small area of the planet’s atmosphere, which cannot exceed the field of view of the last cube acquired (i.e., when the spacecraft is at its closest to Venus), that contains features that can be tracked for the entirety of the sequence. [28] We first identify features that persist throughout the series of observations. That is, we limit our analysis to features that persist for at least 5 h, and do not stray beyond the instrumental field of view. The zonal superrotation
Table 2. Data Analyzed in Evolution Analyses File Name
Orbit
VI0383_00 VI0383_01 VI0383_02 VI0383_03 VI0383_04 VI0383_05 VI0383_06 VI0383_07 VI0383_08 VI0383_09 VI0383_10
383 383 383 383 383 383 383 383 383 383 383
Start Date 8 8 8 8 8 8 8 8 9 9 9
May May May May May May May May May May May
2007 2007 2007 2007 2007 2007 2007 2007 2007 2007 2007
Start Timea
Exposure Timeb (s)
Distance (km)
Resolution (km pixel1)
2028:12 2058:12 2128:12 2158:12 2228:12 2258:12 2328:12 2358:12 0028:12 0058:12 0128:12
3.3 3.3 3.3 3.3 3.3 3.3 3.3 3.3 3.3 3.3 3.3
67106.1 66178.0 65055.9 63735.1 62426.3 60678.6 58710.8 56515.1 54343.1 51631.1 48668.4
16.8 16.5 16.3 15.9 15.6 15.2 14.7 14.1 13.6 12.9 12.2
a
Times and dates are UTC. Exposure time is per slice.
b
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Figure 3. All 10 images considered in the morphological analyses, combined into a single image. The atmosphere is assumed to rotate as a solid body, with a planet-circling period of 5.5 days, as in Figure 1. makes it difficult to be certain of the identity of each feature from image to image, so we crudely correct for the motion of the atmosphere by shifting each image longitudinally at a corresponding rate of about 2.73° hr1 (which equates to a 5.5 day period). This adjustment renders many of the features nearly stationary in the sequence of images, thus identification of the same hole in subsequent images was facilitated. The boundary of each feature is determined in the same manner as the feature boundaries in the morphological analysis above (section 2.1). We measure the peak radiance and the contrast between the peak radiance and the background radiance for each feature that we were able to track in this sequence of images.
well as the overall mean and standard deviation of all of the data considered here. The overall mean indicates that the 10 images that we have analyzed for this paper demonstrate the same tendencies that have been noted by previous
3. Results 3.1. Latitudinal Distribution of Cloud Cover [29] Figure 3 demonstrates the approximate spatial coverage of cloudy and clear regions in the data considered in this paper. All of the considered 1.74 mm images are plotted onto a cylindrical projection. As in Figure 1, the atmosphere is assumed to rotate as a solid body with a period of 5.5 days, so each image has been advected longitudinally according to the time it was acquired relative to the first image (orbit 363). Overlapping regions are averaged together. There is possibly a hint of longitudinal hemispherical asymmetry, similar to that observed by Crisp et al. [1991a], whereby longitudes east of 180° in these data are brighter than those west of 180°; but this conclusion might be biased by the existence of a single very large and very bright feature that spans the region from 220° to 300° longitude and 20° to 5° latitude. Although the data cover nearly two full atmospheric rotations, the coverage and consistency seems to be insufficient to determine whether a longitudinal hemispherical asymmetry exists in these data. [30] Figure 4 shows the latitudinal profile of the mean radiance for each of the images considered in this paper as
Figure 4. Latitudinal profile of the mean radiance for each of the 1.74 mm images considered in this paper. The profiles of the various images are plotted as points. The mean of all of the data considered is shown as a solid line, and range of ±1 standard deviation of all of the data considered is shown with a pair of dashed lines. Also shown here, for comparison, is an estimate of the emitted radiance that results from an application of the cloud properties as derived from the Pioneer Venus Lower Cloud Particle Size Spectrometer (LCPS) by Knollenberg and Hunten [1980]. The vertical dotted line is the calculated radiance; the horizontal dotted line is the latitude of the probe’s descent (the probe actually descended in the northern hemisphere, but we assume roughly hemispherical symmetry only for the purpose of making this comparison).
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Figure 5. Aspect ratio (long axis divided by short axis) as a function of latitude. The data also have been binned into 20° latitude regions, with error bars indicating the standard deviation of aspect ratio in each region. One feature was identified poleward of 60°, which is shown in Figure 5, but obviously could not be sorted into any of the defined latitude bins. observers [Crisp et al., 1991a; Chanover et al., 1998]. Low latitudes tend to be fairly bright, as well as middle to high latitudes (between about 40° to 60°). There is typically less radiance emitted in middle to low latitudes (between about 20° and 40°). The polar collar, which is not easily detected in ground-based observations, can be seen in these images as a region of significantly diminished radiance poleward of about 60°. [31] Also shown in Figure 4, for comparison, is the estimated radiance at the location of the Pioneer Venus Large Probe. Although the Large Probe descended in the northern hemisphere, the location is plotted, in Figure 4, at about 5° latitude. We assume a north-south atmospheric symmetry here because the planetary obliquity is only 3°, indicating minimal seasonal variations, and because previous observations are consistent with such a north-south symmetry, with only a few, probably transient, exceptions. Recent observations by Venus Express indicate that the dynamics of the south polar region are similar to those observed of the north polar region by Pioneer Venus, further supporting the likelihood of north-south hemispheric symmetry [Piccioni et al., 2007]. This calculation of radiance experienced by the Pioneer Venus Large Probe assumes the cloud properties derived by Knollenberg and Hunten [1980], along with a factor of three reduction in the number of mode 3 particles [Pollack et al., 1980]. This represents the radiance that results from a typical assumed set of Venus cloud properties and mass. The optical depth of the middle cloud (i.e., the optical depth below about 57 km) that was used to generate this radiance is approximately t 23. This radiance is calculated using the model of McGouldrick and Toon [2007, 2008a, 2008b]. For a lack of an alternative, previous studies have taken this in situ observation (or reanalyses thereof) of the Venus cloud to be representative of the planetary cloud cover, as a whole [Toon et al., 1982; Grinspoon et al., 1993; James et al., 1997; Imamura and
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Hashimoto, 1998; McGouldrick and Toon, 2007]. Figure 4 shows that this clearly is not the case, since the predicted radiance of the cloud observed in situ by the Pioneer Venus Large Probe is closer to that of the maximum cloudiness (minimum radiance) encountered in the course of the 10 orbits analyzed here. Furthermore, the significant variability in radiance demonstrates the necessity of taking into account both the spatial and temporal cloud variability when determining the radiative balance of the middle atmosphere of Venus. [32] As mentioned in section 1.2, Imamura and Hashimoto [1998] explained that the latitudinal brightness variations seen by Crisp et al. [1991a] likely were a consequence of the global-scale Hadley cell that exists in the atmosphere of Venus. The overall average of the latitudinal distribution of radiance in these data (solid line in Figure 4) is remarkably similar to the results of those simulations. The bright zone around 50° in the data corresponds with the minimum of opacity located around 45° in the simulations of Imamura and Hashimoto [1998]. The diminished radiance poleward of about 60° may be reaching a minimum around 75°. The location of this radiance minimum also corresponds well with the increase in opacity in the nominal simulations of Imamura and Hashimoto [1998], which peak around 70° to 75°. Imamura and Hashimoto [1998] also performed a simulation in which they reduced the effectiveness of the Hadley circulation by reducing by half the speed of the meridional circulation in their simulations. In contrast with the nominal simulation, this second simulation exhibited a peak of cloud opacity somewhat more equatorward, at about 60° latitude (compared with about 75° in their nominal simulation). Although the data considered here represent only 10 days worth of observation, these data indicate that, compared with their test case with reduced meridional speeds, the nominal case of Imamura and Hashimoto [1998] is a better representation of the Hadley circulation of Venus. The increase in the variability of the observed radiance at low latitudes is consistent with the observations by Crisp et al. [1991a] and also suggests the existence of significant convective activity in the clouds of Venus. 3.2. Latitudinal Distribution of Cloud Morphology [33] In Figure 5, we plot the aspect ratio that we measured for approximately 130 features in the 1.74 mm VIRTIS images from orbits 363 through 372, as a function of the latitude of the center of the feature. In addition to plotting each individually measured aspect ratio as symbols, we also bin the measured data into 20° latitude bins. The mean and standard deviation of these three bins are also plotted in Figure 5. We see in Figure 5 that the aspect ratio of the measured holes in the clouds increases with latitude. The typical hole becomes more elongated poleward of 40° (of the 16 features with aspect ratios greater than 5:1, all but three are poleward of 40°). At low latitudes, in addition to holes having aspect ratios closer to 1.0, there also are several features with aspect ratios less than one (i.e., longer in the transverse direction than in the longitudinal direction), something that does not at all occur poleward of 40° in the images considered here. Thus, we observe a preponderance of highly elongated holes at high latitudes, and a
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it will have affected the average brightness in this latitude bin, but not the variability.
Figure 6. Tilt angle (relative to parallels of latitude) as a function of latitude. The data also have been binned into 20° latitude regions, with error bars indicating the standard deviation of tilt angle in each region. An angle of zero means zonally oriented features. The one feature poleward of 60° was not sorted into any bin. tendency for more equal aspect ratios at lower latitudes. These results are consistent with the observations reported by Belton et al. [1991], as summarized in section 1.1. [34] In Figure 6, we plot the angle that each feature makes relative to the parallels of latitude, as a function of latitude. A small ‘‘tilt angle’’ indicates that a feature is aligned nearly zonally, whereas a large tilt angle indicates a feature whose long axis is more vertically oriented. As in Figure 5, we plot each feature individually, as well as the mean and standard deviation of the tilt angle in 20° latitude bins. In Figure 6, we see that the average tilt angle is somewhat larger at low latitudes, becoming smaller at higher latitudes. This, in combination with the trends in Figure 5, indicates that the features not only become more elongated at higher latitudes, they also become more zonally oriented. There is significantly greater variation in the tilt angle at latitudes equatorward of 40°. The smaller variability poleward of 40° suggests that the polar vortex exerts its influence mainly on the poleward side of 40°. Furthermore, the variability equatorward of 40° suggests a complex interaction, likely involving convection, between the local atmospheric dynamics and the formation and evolution of the lower and middle clouds at lower latitudes. [35] Figure 7 shows the peak radiance emerging from holes as a function of latitude. As in Figures 5 and 6, we plot both the individual features themselves, as well as the averages binned over 20° latitude bins. The mean of the feature peak radiance is about 0.2 W m2 mm1 sr1, fairly consistent across all three latitude bins considered here. However, there is significantly more variability in the peak brightness of holes at low latitudes, as compared with the variability seen at higher latitudes. This too suggests that lower latitudes are demonstrating the effects of convection. Since the lower latitudes are viewed most obliquely in these data, the lower latitudes will be most affected by any errors that may occur in the correction for emission angle. However, even if our correction for the emission angle is flawed,
3.3. Feature Lifetimes [36] We might also be able to obtain information about the ages (and, ultimately, the lifetimes) of holes in the clouds from the morphology analyses. Recall, from equation (1), that the change in the atmospheric rotation period with latitude is a function of latitude only, if the zonal velocity is approximately constant. Cloud-tracking analysis by Sa´nchez-Lavega et al. [2008] suggests that the zonal wind is indeed a relatively constant 60 m s1 at latitudes equatorward of about 50°. If we assume that a hole is circular when it is formed, and the zonal velocity is constant, then the poleward edge of the hole advances farther than the equatorward edge of the hole, at a rate given by equation (1) times the latitudinal extent of the hole (Df). If we assume that there is no meridional motion, then the ratio of this translation in degrees longitude to the latitudinal extent of the hole is precisely our previous definition of the tilt angle of the hole! For example, a circular hole centered at 50° latitude, having a meridional extent of 2°, exhibits a roughly 3.15° longitude day1 difference between its top and bottom edges. After 1 day, this equates to a tilt angle of q = tan1(2/3.15) = 32°. After two days, the tilt angle would be 18°, and after three days, 12°. Granted, this is an extreme case. Holes that are closer to the equator experience less shearing because of this effect, since the angular velocity changes less quickly with latitude. A corresponding effect on the aspect ratio can also be calculated. Thus, these measurements of the aspect ratio and of the tilt angle, coupled with knowledge of the horizontal winds in the atmosphere, can provide a rough estimate of the ages of the holes in the clouds. [37] Using the procedure outlined in section 2, we have identified 10 holes, in the series of images taken during orbit 383, whose evolution we have been able to analyze for this paper. This is such a small number of holes that it will be difficult to determine any global trends of hole evolution,
Figure 7. Peak radiance of the holes as a function of feature latitude. The data also have been binned into 20° latitude regions, with error bars indicating the standard deviation of feature peak radiance in each region. The one feature poleward of 60° was not sorted into any bin.
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Figure 8. Derived evolutionary time scales versus peak radiance in the holes. Both axes are parameters of a log linear fit to the measured data. The error bars indicate the standard deviation error estimates of the fitted parameters. whether as a function of position or size. Furthermore, the difficulty in establishing a method of hole definition that is consistent across successive images (i.e., that identifies the same hole, and is not fooled by changes in morphology or surroundings) may introduce small additional uncertainties. For the present investigation, we quantify these numerous error sources in the uncertainties of the fits to the data, as described in subsequent paragraphs. In future investigations we will improve the level of quantification of the error in these analyses of hole evolution, as well as improve our method of feature identification and tracking. Nevertheless, as the first quantitative measurement of cloud evolution, this analysis provides the first benchmark of the lifetimes of the holes in the clouds of Venus. [38] We fit the sequence of the measured peak radiance for each hole according to an exponential variation: ln Ipeak ðt Þ ¼ ln Ipeak ð0Þ
1 t t
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fit to the curve at the time of the first image acquisition. It does not represent an actual measurement of radiance, nor does the variance shown indicate a known measurement error. Figure 8 may exhibit a peak in the evolutionary time constant (i.e., both thicker and thinner holes than this tend to evolve more quickly) at about Ipeak 0.15 W m2 mm1 sr1. However, the small number of holes that have been analyzed, and the large error bars on some of the features, cause us to have only a little confidence in the existence of such a peak, at this time. [40] For Figure 9 we replace Ipeak in equation (2) with the contrast between the peak radiance and the background radiance. We define the contrast as the ratio between the peak radiance of the hole (which we have measured directly from the images), and the local background radiance. We determine the local background radiance to be the mean minus one standard deviation of the radiance in the image at the latitude of the center of the hole. Recall from Figure 4 that the mean minus one standard deviation as a function of latitude tracks remarkably close to the minimum radiance as a function of latitude. Thus, the mean minus one standard deviation stands as a reasonable proxy for the minimum radiance (or thickest cloud cover). In Figure 9, we see a possible trend indicating that holes with greater contrast evolve more quickly than holes with smaller contrast. This is consistent with simulations of hole evolution [McGouldrick and Toon, 2008b]. It also suggests that a uniform cloud cover is the equilibrium situation in the Venus atmosphere, since large contrasts between cloudy and clear regions are quickly reduced, and small contrasts evolve rather slowly. [41] In Figure 10, we plot the evolution time constant obtained from the analysis of the peak radiance as a function of the latitude of the analyzed holes. In Figure 10, the horizontal error bars indicate the range of latitudes at which a given measured hole was located during the course of the image sequence. That is, a large horizontal error bar may be indicative of a significant meridional motion. If meridional
ð2Þ
where the initial peak radiance [ln Ipeak(0)] and the evolutionary time constant (t) are free parameters to be fit to the data for each hole. We find that, for the holes we have analyzed during orbit 383, the average time scale for evolution of a typical hole is 0.95+4.88 0.51 days. Thus, since the typical e-folding time scale of a hole is approximately 1 day, the peak radiance of a typical hole will decrease to approximately 1% of its initial value in a time of about 4.5 days. Note, however, that there is a large upper error bar to this time scale, indicating that a significant fraction of holes can be expected to exhibit time constants for evolution as large as 5 – 6 days. Such slowly evolving holes will remain evident for many weeks. [39] Figure 8 shows the relationship between the initial peak radiance and the evolutionary time scale. Each point in Figure 8 represents the best linear fit to equation (2) for each measured hole. The vertical error bars indicate the one standard deviation variation in the slope (1/t), and the horizontal error bars indicate the one standard deviation variation in the initial peak radiance [Ipeak(0)]. Note that the initial peak radiance in this case refers to the peak radiance
Figure 9. Derived evolutionary time scales versus contrast between the peak hole radiances and the background radiances. Background radiance is defined in the text. Both axes are parameters of a log linear fit to the measured data. The error bars indicate the standard deviation error estimates of the fitted parameters.
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Figure 10. Derived evolutionary time scales versus the latitudes of the measured holes. The horizontal error bars indicate the range of latitudes measured for each hole during the sequence of images. The vertical error bars indicate the standard deviation error estimates of the fitted time scale. motion can be ruled out, then a significant variation in measured feature latitude suggests that we may not be measuring the same feature in each image. With sufficient statistics and independent means of determining wind velocities, such an analysis can provide a further check on the validity of the analysis of feature evolution. Even with the small number of features measured in this instance, there appears to be a trend in which holes at lower latitudes evolve more quickly than those at higher latitudes. This behavior is consistent with the earlier assessment of increased cloud variability at equatorial latitudes. [42] In each of these analyses, our conclusions are necessarily tentative, as they derive from measurements of only 10 holes over the span of only 5 h. We hope to improve on these statistics by measuring many more features in many more images in the future. Nevertheless, these analyses represent a first quantitative assessment of the evolution of holes in the middle and lower clouds of Venus.
4. Conclusions [43] We have analyzed a series of ten 1.74 mm VIRTIS-M IR images, taken on successive orbits, covering a span of 10 days, to identify tendencies to the morphology and the brightness of the holes in the clouds of Venus. We have also analyzed a separate series of eleven 1.74 mm VIRTIS-M IR images, taken over the course of a single orbit, spanning a time of 5 h, to investigate the evolution of the holes in the clouds of Venus. [44] We find that the latitudinal distribution of radiance in which midlatitudes are relatively brighter, and low latitudes exhibit somewhat greater overall variability, is consistent with that observed by Crisp et al. [1991a], indicating that this tendency is a fairly stable phenomenon. The latitudinal distribution of average radiance that we observe in the images considered here is largely consistent with the simulations of Imamura and Hashimoto [1998], indicating that the global Hadley circulation plays a significant role in the distribution of the cloud mass of Venus. We observe an
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increase in the variability of radiance at lower latitudes, suggestive of the importance of convection to the nature of the middle and lower clouds of Venus. While the distribution of potential cloud constituents and their precursors (i.e., H2SO4, SO2, H2O, and amorphous sulfur) are not constrained by the data and analyses presented here, variations in these constituents, driven either by chemistry or by the atmospheric dynamics already noted, might also play a role in the cloud variability. [45] We find that the morphology of the holes in the clouds of Venus trends from features having a wide variety of orientations (relative to the zonal axis) and aspect ratios of approximately 1.0 at lower latitudes, to highly elongated and zonally oriented features at high latitudes. The greater variability in the morphology of features at lower latitudes is possibly indicative of a greater amount of convective activity there. We fail to find a correlation between the peak radiance of a hole and its latitude. A hint of a longitudinal hemispheric asymmetry, similar to that observed by Crisp et al. [1991a] and Chanover et al. [1998] was seen; but the amount of data considered here is insufficient to be certain of its existence here. [46] We find that, if the data analyzed here are typical, holes in the middle and lower clouds of Venus evolve with a time constant of approximately 1 day. We find that holes exhibiting greater contrast between their peak radiance and the background radiance, and holes at lower latitudes evolve more quickly. This is consistent with the increased variability seen at lower latitudes that results from enhanced convection. It is also consistent with the notion that the middle and lower clouds are supported by a radiative dynamical feedback, whereby regions exhibiting large contrasts in clouds cover will also experience large contrasts in radiative heating because of the absorption of upwelling infrared radiation. However, the data considered here represent only 5 h of observation of a fairly small area (about 2% of the atmosphere). This is hardly sufficient data to allow us to draw any broad conclusions. We intend to carry out future efforts to improve both the methods of the analysis as well as the volume of analyzed data. [47] The features in the images analyzed here strongly indicate the activity of convection in the lower and middle cloud decks of Venus. The spatial variability and short time scale for their evolution are somewhat inconsistent with typical terrestrial stratus, to which the clouds of Venus are often compared. However, previous radiative modeling by Pollack et al. [1980] and microphysical modeling of the middle and lower clouds by McGouldrick and Toon [2007] suggests the existence of a radiative dynamical feedback, which can support convection in the lower and middle cloud decks of Venus. The clouds of Venus may be more analogous to terrestrial stratocumulus ‘‘cloud streets,’’ whereby convection exists in the midst of strong horizontal flow [Houze, 1993]. Depending upon the relative strengths of the horizontal flow and the convection, either rolls or cellular patterns may be formed. Similarly, the nature of terrestrial thunderstorm development (whether single-cell, multicell, or supercell) is known to be somewhat dependent upon the local wind shear and static stability [Houze, 1993]. Perhaps processes analogous to cloud street formation or the development of multicell thunderstorms are occurring in the atmosphere of Venus in the vicinity of the lower and middle clouds.
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[48] Here we have presented analyses of only a small subset of the VIRTIS data. We have analyzed data from only 11 orbits, from among more than 500: barely 2% of the total available VIRTIS data set. However, given that at this writing, Venus Express has been orbiting Venus for more than three full Venus years, there likely is sufficient data in the archive to seek out potential tendencies of hole formation and/or appearance as a function of geographic or solar longitude. We intend to carry out further analyses of the VIRTIS data to determine whether such tendencies exist. Furthermore, comparisons with observations at other wavelengths, and concurrent Venus Monitoring Camera observations, can be utilized to gain greater understanding of the properties and evolution of the Venus clouds. In the longer term, the anticipated arrival of the Japanese Space Agency’s Venus Climate Orbiter (VCO) in 2010 promises to enhance the study of the Venus clouds [Nakamura et al., 2007]. The understanding of the nature of the middle and lower clouds of Venus could gain much from the potential coordinated observation of Venus by the two spacecraft. Subsequent observations of the same region of the atmosphere could extend the analysis of cloud evolution beyond the limits that either VCO or Venus Express could accomplish by itself, by extending the baseline of time over which a particular feature is viewable. Concurrent observations of features by multiple spacecraft could further elucidate the cloud properties. For example, it may be possible to determine the difference between the properties of a profile of a hole versus that of a cloud by performing occultation measurements with one spacecraft, while the other identifies the types of features that are passing over the limb of the planet. Alternately, long-lived balloon observatories can be placed in the atmosphere and possibly maneuvered to sample the atmosphere in regions that are identified by the concurrently orbiting spacecraft. [49] Acknowledgments. We thank two anonymous reviewers for their careful reading of the manuscript and their suggestions for making this a stronger paper. We also acknowledge conversations with Colin Wilson, Con Tsang, and Robert Carlson that have helped to improve the analysis of the data with respect to the correction for the emission angle. We also thank The European Space Agency for the opportunity to work with Venus Express. We thank Pierre Drossart, Giuseppi Piccioni, and the VIRTIS team for their assistance in enabling us to reduce and analyze the VIRTIS data. A portion of the work described in this paper was carried out at the Jet Propulsion Laboratory, Pasadena, California, under contract with NASA. K.M. and D.H.G. were supported by NASA in support of ESA’s Venus Express mission under grant NNX07AI61G.
References Allen, D. A., and J. W. Crawford (1984), Cloud structure on the dark side of Venus, Nature, 307, 222 – 224. Baker, R. D., G. Schubert, and P. W. Jones (2000), Convectively generated internal gravity waves in the lower atmosphere of Venus, part II: Mean wind shear and wave-mean flow interaction, J. Atmos. Sci., 57, 200 – 215. Belton, M. J. S., et al. (1991), Images from Galileo of the Venus cloud deck, Science, 253, 1531 – 1536. Blamont, J. E., et al. (1986), Implications of the VEGA balloon results for Venus atmospheric dynamics, Science, 231, 1422 – 1425. Bullock, M. A., and D. H. Grinspoon (2001), The recent evolution of climate on Venus, Icarus, 150, 19 – 37, doi:10.1006/icar.2000.6570. Carlson, R. W., et al. (1991), Galileo infrared imaging spectroscopy measurements at Venus, Science, 253, 1541 – 1548. Carlson, R. W., L. W. Kamp, K. H. Baines, J. B. Pollack, D. H. Grinspoon, Th. Encrenaz, P. Drossart, and F. W. Taylor (1993), Variations in Venus cloud particle properties: A new view of Venus’s cloud morphology as observed by Galileo Near-Infrared Mapping Spectrometer, Planet. Space Sci., 41, 477 – 485, doi:10.1016/00320633(93)900306.
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Chanover, N. J., D. A. Glenar, and J. J. Hillman (1998), Multispectral nearIR imaging of Venus nightside cloud features, J. Geophys. Res., 103, 31,335 – 31,348. Counselman, C. C., S. A. Gourevitch, R. W. King, G. B. Loroit, and E. S. Ginsberg (1980), Zonal and meridional circulation of the lower atmosphere of Venus determined by radio interferometry, J. Geophys. Res., 85, 8026 – 8030. Crisp, D. (1989), Radiative forcing of the Venus mesosphere, Icarus, 77, 391 – 413, doi:10.1016/00191035(89)900961. Crisp, D., et al. (1989), The nature of the near-infrared features on the Venus night side, Science, 246, 506 – 509. Crisp, D., D. A. Allen, D. H. Grinspoon, and J. B. Pollack (1991a), The dark side of Venus: Near-infrared images and spectra taken from the Anglo-Australian Observatory, Science, 253, 1263 – 1266. Crisp, D., et al. (1991b), Ground-based near-infrared imaging observations of Venus during the Galileo encounter, Science, 253, 1538 – 1541. Drossart, P., et al. (2007), Scientific goals for the observation of Venus by VIIS on ESA/Venus Express mission, Planet. Space Sci., 55, 1653 – 1672, doi:10.1016/j.pss.2007.01.003. Grinspoon, D. H., J. B. Pollack, B. R. Sitton, R. W. Carlson, L. W. Kamp, K. H. Baines, Th. Encrenaz, and F. W. Taylor (1993), Probing Venus’s cloud structure with Galileo NIMS, Planet. Space Sci., 50, 515 – 542, doi:10.1016/00320633(93)90034Y. Hashimoto, G. L., and Y. Abe (2000), Stabilization of Venus’ climate by a chemical-albedo feedback, Earth Planets Space, 52, 197 – 202. Hinson, D. P., and J. M. Jenkins (1995), Magellan radio occultation measurements of atmospheric waves on Venus, Icarus, 114, 310 – 327, doi:10.1006/icar.1995.1064. Houze, R. A. (1993), Cloud Dynamics, Academic, San Diego, Calif. Imamura, T., and G. L. Hashimoto (1998), Venus cloud formation in the meridional circulation, J. Geophys. Res., 103, 31,349 – 31,366. James, E. P., O. B. Toon, and G. Schubert (1997), A numerical microphysical model of the condensational Venus cloud, Icarus, 129, 147 – 171, doi:10.1006/icar.1997.5763. Knollenberg, R. G., and D. H. Hunten (1980), The microphysics of the clouds of Venus: Results of the Pioneer Venus Particle Size Spectrometer Experiment, J. Geophys. Res., 85, 8038 – 8058. Linkin, V. M., et al. (1986), Vega balloon dynamics and vertical winds in the Venus middle cloud region, Science, 231, 1417 – 1419. Markiewicz, W., et al. (2007a), Venus Monitoring Camera for Venus Express, Planet. Space Sci., 55, 1701 – 1711, doi:10.1016/j.pss.2007.01.004. Markiewicz, W. J., et al. (2007b), Morphology and dynamics of the upper cloud layer of Venus, Nature, 450, 633 – 636, doi:10.1038/nature06320. McGouldrick, K., and O. B. Toon (2007), Investigation of possible causes of the holes in the condensational Venus cloud using a microphysical cloud model with a radiative-dynamical feedback, Icarus, 191, 1 – 24, doi:10.1016/j.icarus.2007.04.007. McGouldrick, K., and O. B. Toon (2008a), Observable effects of convection and gravity waves on the Venus condensational cloud, Planet. Space Sci., 46, 1112 – 1131, doi:10.1016/j.pss.2008.02.010. McGouldrick, K., and O. B. Toon (2008b), Modeling the effects of shear on the evolution of the holes in the condensational clouds of Venus, Icarus, 196, 35 – 48, doi:10.1016/j.icarus.2008.02.020. Nakamura, M., et al. (2007), Planet-C: Venus Climate Orbiter mission of Japan, Planet. Space Sci., 55, 1831 – 1842, doi:10.1016/j.pss.2007.01.009. Piccioni, G., et al. (2007), South-polar features on Venus similar to those near the north pole, Nature, 450, 637 – 640, doi:10.1038/nature06209. Pollack, J. B., O. B. Toon, and R. Boese (1980), Greenhouse models of Venus’ high surface temperature, as constrained by Pioneer Venus measurements, J. Geophys. Res., 85, 8223 – 8231. Sa´nchez-Lavega, A., et al. (2008), Variable winds on Venus mapped in three dimensions, Geophys. Res. Lett., 35, L13204, doi:10.1029/ 2008GL033817. Schubert, G., et al. (1980), Structure and circulation of the Venus atmosphere, J. Geophys. Res., 85, 8007 – 8025. Taylor, F. W., et al. (1979), Temperature, cloud structure, and dynamics of Venus middle atmosphere by infrared remote-sensing from pioneer orbiter, Science, 205, 65 – 67. Toon, O. B., R. P. Turco, and J. B. Pollack (1982), The ultraviolet absorber on Venus: Amorphous sulfur, Icarus, 51, 358 – 373, doi:10.1016/ 00191035(82)900896.
K. H. Baines and T. W. Momary, Jet Propulsion Laboratory, California Institute of Technology, M/S 183-601, 4800 Oak Grove Drive, Pasadena, CA 91109, USA. D. H. Grinspoon and K. McGouldrick, Department of Space Sciences, Denver Museum of Nature and Science, 2001 Colorado Boulevard, Denver, CO 80205, USA. (
[email protected])
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Whistler mode waves from lightning on Venus: Magnetic control of ionospheric access C. T. Russell,1 T. L. Zhang,2 and H. Y. Wei1 Received 29 February 2008; revised 29 May 2008; accepted 19 June 2008; published 18 September 2008.
[1] The fluxgate magnetometer on Venus Express samples the magnetic field near
periapsis at 128 Hz. Bursts of plane-polarized magnetic waves in the vicinity of 100 Hz are observed propagating at small angles to the magnetic field. The magnetic field is generally horizontal in the region around periapsis, located at high northern latitudes. When the magnetic field remains within 15° of horizontal during the 2-min periapsis pass, no such waves are observed; but when there are brief periods during which the local magnetic field dips into the atmosphere by more than 15°, the bursts begin to appear. Such radial excursions of the magnetic field occur 25% of the time in the region around periapsis. The bursts are seen only on passes with these excursions. We interpret this magnetic control in terms of the coupling between the electromagnetic wave from lightning discharges refracted vertically by the increasing electron density and the nearly horizontal ionospheric magnetic field along which the energy is guided to the spacecraft. The inferred rate of electric discharges in the Venus atmosphere is about 20% of that seen in the Earth’s atmosphere. Citation: Russell, C. T., T. L. Zhang, and H. Y. Wei (2008), Whistler mode waves from lightning on Venus: Magnetic control of ionospheric access, J. Geophys. Res., 113, E00B05, doi:10.1029/2008JE003137.
1. Introduction [2] While the Venus atmosphere is quite dry, containing very little water, Venus is also quite cloudy, with hydrated sulfuric acid droplets shrouding the planet some 50 –60 km above the surface. At low altitudes, the winds are slow, but in the clouds, the horizontal velocities are observed to range to over 100 ms 1. These ‘‘4-d’’ winds produce much wind shear and probably are accompanied by significant vertical transport as well. Thus, the clouds may well be electrified as terrestrial clouds are. The observations reported herein are over the polar vortex that is possibly associated with rapid downwelling [Piccioni et al., 2007]. [3] Many researchers have reported optical and electromagnetic signals that could be produced by discharges in or from these clouds. Krasnopolsky [1983] reported flashes seen in the Venera-9 visible spectrometer; Hansell et al. [1995] reported detections with an Earth-based telescope. The landing probes of Venera 11, 12, 13, and 14 detected electromagnetic pulses as they descended through the atmosphere and as they sat on the surface [Ksanfomaliti, 1983]. They have been most extensively studied from orbit on Pioneer Venus (PVO) using an electric antenna [Taylor et al., 1979; Scarf et al., 1980]. Finally, at radio frequencies that are able to propagate through the ionosphere with minimal interaction with the plasma, Galileo recorded radio 1 Institute of Geophysics and Planetary Physics, University of California, Los Angeles, California, USA. 2 Space Research Institute, Austrian Academy of Sciences, Graz, Austria.
waves similar to terrestrial radio frequency emissions from lightning [Gurnett et al., 1991]. However, as Gurnett et al. [2001] later reported, these high-frequency radio waves are weaker in strength and less frequent in occurrence than terrestrial lightning. [4] The waves detected in the magnetized Venus ionosphere by PVO were of two distinct types [Russell, 1991]. One class appeared only when the magnetic field had a significant radial component and behaved like an electromagnetic wave would. This wave diminished only slightly in estimated electromagnetic energy flux with increasing altitude. The second class was apparently electrostatic and attenuated rapidly with altitude. It was seen in all four narrowband frequency channels: 100 Hz, 730 Hz, 5.6 kHz, and 30 kHz. We interpret these electrostatic waves as being associated with cloud-to-ionosphere strikes, probably ‘‘local.’’ The electromagnetic waves were most probably due to intracloud discharges, possibly at some distance. [5] The occurrence of lightning on Venus remained controversial after the Venera and PVO missions ceased because some searches were unsuccessful. A search for scattered light in the PVO star sensor came up empty [Borucki et al., 1991], albeit the total time the star sensor was active and able to see flashes was only a few minutes in total, with very little coverage over the region defined as active by the PVO electric field instrument. The photometer on the Venus balloons also did not observe flashes [Sagdeev et al., 1986], but the balloons were in the clouds and were not looking down on the clouds from above. Thus any light from flashes could be scattered before reaching the detectors on the balloons.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003137$09.00
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Figure 1. Waveform of 8-ms samples of the magnetic field at periapsis on 1 July 2006, bandpass filtered from 42 to 60 Hz and rotated into principal axis coordinates.
[6] The waves observed by Pioneer Venus were quite strong. In a vacuum, the energy flux of electromagnetic waves is evenly divided between the electric and magnetic components. In the ionosphere where the index of refraction is high, the wave slows down and the energy is mainly carried by the magnetic component. On Pioneer Venus, the plasma density, the magnetic field, and the electric component of the wave were measured and we could calculate the electromagnetic energy of the waves identified as propagating in the whistler mode [Russell et al., 1989]. Furthermore, we could check this estimate with data obtained in the atmosphere beneath the ionosphere on the last few orbits when Pioneer Venus began to enter the atmosphere because of gravitational perturbations and eventually because of drag on the spacecraft [Strangeway et al., 1993]. These calculations indicated that the Pioneer Venus electric waves near 100 Hz could be detected by a state-of-the-art fluxgate magnetometer in the Venus ionosphere [Russell et al., 2006]. The Pioneer Venus magnetometer did not have the bandwidth to detect these waves. When it was decided to include a magnetometer on Venus Express, the decision was made to include a 128-Hz sampling mode that would allow magnetic waves near 100 Hz to be sampled [Zhang et al., 2006]. These data were obtained initially for only 2 min around periapsis. In late December 2006, this interval was extended. As expected, waves with the expected amplitude and temporal structure were observed at periapsis [Russell et al., 2007]. In this paper we report on the properties of the waves seen in these 2-min sampling intervals.
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has been described in detail by Zhang et al. [2006], and the cleaning process has been described by Zhang et al. [2008]. The Alfvenic character of interplanetary fluctuations has been used to establish the magnetometer’s DC level to better than 1 nT [Leinweber et al., 2008]. [8] In addition to the DC spacecraft fields, there is AC interference, in particular from the reaction wheels. These waves are found in the passband in which we expect to see the waves due to lightning. They also can vary in frequency from pass to pass, and at times, during a pass. Examples of these waves have been published by Russell et al. [2008]. A very simple test of whether a wave is in the ambient plasma or due to spacecraft noise is simply to subtract the measurements of the two magnetometers. Natural waves have the same amplitude at the two locations and disappear in the difference wave. The reaction wheel noise is stronger in the sensor on the spacecraft deck and does not disappear in subtraction [Russell et al., 2008]. As we demonstrate below in section 3, the bursts we discuss in this report pass this test. [9] In this study we use the outbound sensor and pass the time series through a bandpass filter from 42 to 60 Hz. This range of frequencies is chosen to avoid strong signals from the reaction wheels [Russell et al., 2008]. The magnetometer has an antialiasing filter, but this filter does allow strong signals above the Nyquist frequency of 64 Hz to enter the telemetry stream. These signals will appear to be below 64 Hz, but their polarization will be reversed. When reaction wheel noise appears in this passband, we simply eliminate this pass from consideration. We examine herein observations obtained in 2006, after the initial commissioning of the instruments in orbit.
3. Observations [10] In the Pioneer Venus observations, whistler mode waves could not be detected in the day-lit hemisphere because of solar-induced noise in the plasma wave electric antenna. One of the advantages of the Venus Express magnetic measurement is that the waves can be observed in sunlight as well as darkness. Figure 1 shows a filtered
2. Instrument and Data Processing [7] Because Venus Express was a low-cost mission and employed an existing bus, there was no magnetic cleanliness program. Instead, two magnetometer triads were installed: one on the top deck and the second on the end of a 1-m boom. The two magnetometers were aligned and used in a gradiometer mode to identify spacecraft-associated waves so that they could be removed. The magnetometer
Figure 2. Three seconds of the record shown in Figure 1 showing substructure in the burst.
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Figure 3. Hodogram of 400 ms of the burst shown in Figure 2 in the principal axis system.
waveform from the outer fluxgate sensor on 1 July 2006, at 308-km altitude and 0642 LT. The wave packets have been rotated into their principal axis system where the minimum variance direction is along the third axis, and the maximum variance direction is along the first. Clearly, the waves are confined to two directions and have similar amplitudes in these two directions. Figure 2 focuses on 3 s of this record revealing substructure within the burst. Overall, the wave packets last about 1/4 to 1/2 s, but clearly substructure can occur on 100-ms scales. A hodogram of these waves is shown in Figure 3, over 400 ms of data. The waves are confined to a plane as is expected of whistler mode signals. The waves’ phase velocity (k vector) along the minimum variance direction is propagating at only 9° to the background magnetic field. As mentioned above, the magnetometer samples continuously in a gradiometer mode with simultaneous samples at the end of the boom and on the spacecraft upper deck. Ambient waves will be the same
Figure 4. By component of the 128-Hz magnetic field in spacecraft coordinates measured on the outboard and inboard sensors and the difference between the two readings. Naturally occurring, nonspacecraft sources would be identical at the two locations and cancel in subtraction. The burst’s absence indicates that the burst in Figure 3 is not due to spacecraft sources.
Figure 5. Bandpass-filtered measurements in principal axis system for 400 ms at periapsis on 23 December 2006. Comments of Figure 1 apply.
amplitude at the two magnetometers so that subtracting the two time series will result in a time series of only noise. In Figure 4, we demonstrate this subtraction for the event shown in the hodogram of Figure 3. The top trace is the output of the y component outboard sensor. The middle trace is the y component inboard sensor. The bottom trace is the difference. In the difference trace, the burst of waves vanishes into the noise level, indicating it is an ambient signal. [11] An event observed on 23 December 2006, at 284-km altitude, 88° latitude, 1042 LT, and 88.5° solar zenith angle is shown in Figure 5. This event is propagating at only 6° to the local magnetic field. Figure 6 shows a hodogram of 250 ms of data at the maximum of the event. Figure 7 repeats our gradiometer difference test for the burst in the hodogram in Figure 6. Again the burst disappears in the difference trace indicating its source is not on the spacecraft. [12] Figure 8 shows projections of the background magnetic field along the orbit in the plane containing the solar direction and the orbit plane normal. The 8-s averaged magnetic field is shown every minute. The event shown in Figure 7 occurs at the second vector before 0750 UT
Figure 6. Hodogram of 250 ms of the burst shown in Figure 5.
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where the field is dipping most strongly into the atmosphere. This suggests that we should examine the control of the occurrence of these waves by the magnetic field orientation. We demonstrate this control statistically in section 4.
shows that on some days, there just are not waves present near the dipping field lines. July 2 has an ambient magnetic field that is quite horizontal and says little about the possible presence of lightning near the satellite at this time. We note that more reaction wheel noise is present here than on other passes, but we feel we would have seen a typical amplitude burst if present. [16] Figure 11 shows three consecutive days in December 2006, when the spacecraft was near 251 km, 0141 LT, and 96.8° solar zenith angle. December 22 has a horizontal ambient magnetic field and no wave events. December 23 has a field inclined nearly 40° to the horizontal and has a strong wave event, near 1 nT peak-to-peak. December 24 deviates nearly 20° to the horizontal, and no wave bursts are seen. Again we have a pattern of inclined field being associated with occasional wave events and horizontal fields being associated with quiet intervals. [17] Figure 12 summarizes our survey of the 2006 Venus Express data. Plotted is the peak-to-peak amplitude of the strongest bursts on each pass versus the maximum deviation of the ambient field from the horizontal. We see that when the field deviation was less than 15° from the horizontal direction everywhere on the pass, no whistler mode wave bursts were detected. When the deviation was greater than 17°, the wave events could occur. We attribute this magnetic control to the effect of the ionosphere on the electromagnetic waves propagating from the clouds below and the properties of the propagation of whistler mode waves at the conditions in the Venus ionosphere as discussed in section 5.
4. Magnetic Field Control of Burst Occurrence
5. Discussion
[13] To illustrate the nature of the magnetic field control of the appearance of the whistler mode waves, we display the entire 2-min passes for three consecutive orbits together with the angle between the ambient magnetic field and the local horizontal direction for the same interval plotted with 1-s resolution. These orbits have very similar altitudes and local times but may differ in the steadiness and orientation of the magnetic field. [14] Figure 9 shows the measurements on 8, 9, and 10 June 2006. On 10 June, the ambient magnetic field is quiet and close to horizontal for the entire 2-min pass. No wave bursts are seen in the right-hand panel. In contrast on 9 June, the ambient magnetic field is time varying moving up to 30° away from horizontal. Strong wave packets accompany these variations at the beginning of the interval. These wave packets are analyzed in detail by Russell et al. [2007]. June 8 has similar field variations to those of 9 June, but no wave events are seen. This observation indicates that the magnetic orientation is not the only controlling factor. In fact, the most probable explanation of the lack of waves on this pass is that there are no atmospheric discharges below the satellite on this day. [15] Another 3-day interval, 30 June, 1 and 2 July, is shown in Figure 10 in the same format as Figure 9. In this interval, there is a burst on 1 July, but no activity on 30 June or 2 July. The most inclined fields occur on 30 June, but they are not accompanied by whistler mode bursts. The 1 July ambient fields do deviate from the horizontal by about 20° and are sufficient to allow the entry of waves into the ionosphere as in the previous example, but clearly 30 June
[18] The evidence for the necessity of inclined ionospheric magnetic fields for the access of whistler mode
Figure 7. By component of the 128-Hz field in spacecraft coordinates measured on the outboard and inboard sensors and the difference between the two readings. The absence of the burst on the bottom trace indicates that the burst in Figure 6 is not due to spacecraft sources.
Figure 8. Venus Express orbit together with projections of 8-s samples of the magnetic field once per minute in the x-z plane in Venus Solar Orbital coordinates. Magnetic field is seen to be dipping into atmosphere around time of burst of noise.
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Figure 9. Three consecutive periapsis passes on 8, 9, and 10 June. The left-hand panels show the angle of the field to the radial direction with 90° being horizontal as calculated from 1-s samples of the field. The right-hand panels show bandpass-filtered 8-ms magnetic field measurements covering the frequency range from 40 to 62 Hz. The component shown is from the y axis sensor in spacecraft coordinates. This is the sensor with the lowest noise level. waves from the atmosphere into the ionosphere is clear from our statistical survey. However, in order to understand why this access control occurs, we need to examine the expected characteristics of the electromagnetic waves that enter the
ionosphere from below. Figure 13 shows a cartoon of the wavefronts produced by an intracloud discharge in Venus’ atmosphere. The waves at first propagate with spherical wavefronts at the speed of light. When they encounter the
Figure 10. Three consecutive passes on 30 June, 1 July, and 2 July 2006. Comments of Figure 9 apply. 5 of 8
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Figure 11. Three consecutive passes on 22, 23, and 24 December 2006. Comments of Figure 9 apply. ionosphere, the part of the wavefront that enters the ionosphere slows down considerably to a very small fraction of the velocity of light depending on the electron density and the strength of the magnetic field. The flat wavefront corresponds to vertical propagation of the wave. The
Figure 12. Peak-to-peak amplitudes of burst maxima on each pass versus maximum inclination of the magnetic field to the horizontal.
Figure 13. Cartoon of electric discharge in the Venus clouds and the resultant wavefronts of electromagnetic waves. When the wavefronts enter the ionosphere where the phase velocity decreases substantially, the wavefront flattens and the wavenormals become vertical.
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Figure 14. Phase and group velocities of whistler mode waves at 100 Hz propagating at various angles to the magnetic field in a plasma with a density of 5000 cm 3 and a magnetic field strength of 20 nT. magnetic field in the ionosphere, as we have seen, tends to be horizontal so long as the interplanetary magnetic field is steady. Inclined fields can be produced by time variations of the magnetic fields that in turn can lead to interconnection of different field directions within the ionosphere. [19] Figure 14 shows the phase and group velocities of a whistler mode wave at 100 Hz propagating at various angles to the magnetic field in a plasma with a density of 5000 electrons cm 3 and a magnetic field strength of 20 nT. These velocities are calculated from the full dispersion relation for cold plasma waves [Stix, 1962]. On the left is the phase velocity. This wave does not travel perpendicular to the magnetic field. In fact, there is a cone of nonpropagation of the order of 10°, an angle similar to the angle of inclination needed for the waves to reach Venus Express. On the right is the group velocity in a similar display. Once the energy couples to the magnetic field, it is strongly guided by the magnetic field. Thus, our observations are consistent with what one would expect for the waves in the ionosphere that would be produced as a result of atmospheric electrical discharges. [20] We can use the first year’s statistics of the magnetic field orientation and the number of discrete whistler wave bursts during the 2-min sampling intervals to make a rough comparison with terrestrial rates. If we use the conservative assumption that access to the ionosphere occurs only when the field is inclined 15° or more to the horizontal, then the statistics of cone angle occurrence indicates that access is allowed only 25% of the time. In the first Venus year of operation up to 16 December 2006, we received 12,223 s of 128-Hz data that could be analyzed. During this interval, 61 individual bursts of noise were detected, giving a rate of 0.005 per second. Normalizing this rate for access time, i.e., the duration of the fields inclined more than 15° to the horizontal, we get a rate of 0.02 per second below the clouds. If the spacecraft can see a circle of 200-km radius in the clouds below it, a distance equal to its height above them, it is seeing only 0.027% of the surface. If we use this ratio to normalize our observations, we get a global rate of 18 strokes s 1 or 20% the terrestrial rate. We do not know if our region of study near Venus’ polar vortex is representative of the entire planet or if our area of seeing is correct, but these numbers do tell us that lightning is a significant and important process in the Venus atmosphere. Finally, we note
that our measurements were gathered over a full Venus year covering all local times. While we would not necessarily expect the results of our observations to agree with the rate seen on a brief Cassini swing-by, observing radio waves and not the whistler mode waves studied here, we do agree qualitatively with the result of Gurnett et al. [2001] that the rate at Venus may be lower than on Earth.
6. Summary [21] In summary, Venus Express observes bursts of whistler mode noise in the Venus ionosphere. These waves are almost circularly plane-polarized and propagate at a small angle to the magnetic field. They were seen only when the local ionospheric magnetic field was inclined 17° or more to the horizontal. Such a cone of evanescence would be expected for electromagnetic radiation produced in the atmosphere and refracted vertically as it enters the ionosphere. The observations are consistent with lightning being prevalent in the Venus atmosphere and creating the whistler mode waves seen in the ionosphere, both those observed with the electric antenna on Pioneer Venus at low latitudes and those detected with the Venus Express magnetometer at high latitudes. [22] Acknowledgments. This work was supported by the National Aeronautics and Space Administration under research grant NNG06GC62G.
References Borucki, W. J., et al. (1991), Pioneer Venus Orbiter search for Venusian lightning, J. Geophys. Res., 96(A7), 11,033 – 11,043, doi:10.1029/ 91JA01097. Gurnett, D. A., et al. (1991), Lightning and plasma wave observations from the Galileo fly by of Venus, Science, 253(5027), 1522 – 1525, doi:10.1126/science.253.5027.1522. Gurnett, D. A., et al. (2001), Non-detection at Venus of high-frequency radio signals characteristic of terrestrial lightning, Nature, 409(6818), 313 – 315, doi:10.1038/35053009. Hansell, S. A. W. K. W., et al. (1995), Optical detection of lightning on Venus, Icarus, 117, 345 – 351, doi:10.1006/icar.1995.1160. Krasnopolsky, V. A. (1983), Lightnings and nitric oxide on Venus, Planet. Space Sci., 31, 1363 – 1369, doi:10.1016/0032-0633(83)90072-7. Ksanfomaliti, L. V. (1983), Electrical activity of the atmosphere of Venus: I. Measurements on descending probes, Kosm. Issled., 21, 279 – 296. Leinweber, H., C. T. Russell, K. Torkar, J. L. Zhang, and V. Angelopoulos (2008), An advanced approach to finding magnetometer zero levels in the interplanetary magnetic field, Meas. Sci. Technol., 19, doi:10.1088/09570233/19/5/055104.
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Piccioni, G., et al. (2007), South-pole features on Venus similar to those near the north pole, Nature, 450(7170), 637– 640, doi:10.1038/nature06209. Russell, C. T. (1991), Venus lightning, Space Sci. Rev., 55, 317 – 356. Russell, C. T., et al. (1989), VLF bursts in the night ionosphere of Venus: Estimates of the Poynting flux, Geophys. Res. Lett., 16(6), 579 – 582, doi:10.1029/GL016i006p00579. Russell, C. T., R. J. Strangeway, and T. L. Zhang (2006), Lightning detection on the Venus Express Mission, Planet. Space Sci., 54, 1344 – 1351, doi:10.1016/j.pss.2006.04.026. Russell, C. T., T. L. Zhang, M. Delva, W. Magnes, R. J. Strangeway, and H. Y. Wei (2007), Lightning on Venus inferred from whistler-mode waves in the ionosphere, Nature, 450(7170), 661 – 662, doi:10.1038/ nature05930. Russell, C. T., T. L. Zhang, R. J. Strangeway, H. Y. Wei, M. Delva, and W. Magnes (2008), Electromagnetic waves observed by Venus Express at periapsis: Detection and analysis techniques, Adv. Space Res., 41(1), 113 – 117, doi:10.1016/j.asr.2007.08.032. Sagdeev, R. Z., et al. (1986), Overview of VEGA Venus balloon in-situ meteorological measurements, Science, 231(4744), 1411 – 1414, doi:10.1126/science.231.4744.1411. Scarf, F. L., et al. (1980), Lightning on Venus: Orbiter detection of whistler signals, J. Geophys. Res., 85(A13), 8158 – 8166, doi:10.1029/ JA085iA13p08158.
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Stix, T. H. (1962), Theory of Plasma Waves, 283 pp., McGraw-Hill, New York. Strangeway, R. J., et al. (1993), Observation of intense wave bursts at very low altitudes within the Venus nightside ionosphere, Geophys. Res. Lett., 20(23), 2771 – 2774, doi:10.1029/93GL02702. Taylor, W. W. L., et al. (1979), Absorption of whistler mode waves in the ionosphere of Venus, Science, 205(4401), 112 – 114, doi:10.1126/ science.205.4401.112. Zhang, T. L., et al. (2006), Magnetic field investigation of the Venus plasma environment: Expected new results from Venus Express, Planet. Space Sci., 54, 1336 – 1343, doi:10.1016/j.pss.2006.04.018. Zhang, T. L., et al. (2008), Initial Venus Express magnetic field observations of the Venus bow shock location at solar minimum, Planet. Space Sci., 56, 785 – 789, doi:10.1016/j.pss.2007.09.012. C. T. Russell and H. Y. Wei, Institute of Geophysics and Planetary Physics, University of California, 603 Charles Young Drive East, 3845 Slichter Hall, Los Angeles, CA 90095-1567, USA. (
[email protected]. edu) T. L. Zhang, Space Research Institute, Austrian Academy of Sciences, Schmiedlstrasse 6, A-8042 Graz, Austria.
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Full Article
Morphology and dynamics of Venus oxygen airglow from Venus Express/Visible and Infrared Thermal Imaging Spectrometer observations R. Hueso,1 A. Sa´nchez-Lavega,1 G. Piccioni,2 P. Drossart,3 J. C. Ge´rard,4 I. Khatuntsev,5 L. Zasova,5 and A. Migliorini2 Received 21 January 2008; revised 21 March 2008; accepted 2 May 2008; published 26 July 2008.
[1] Images obtained by the Visible and Infrared Thermal Imaging Spectrometer
(VIRTIS)-M channel instrument onboard Venus Express have been used to retrieve maps and apparent motions of the O2 (1D) infrared nightglow on Venus at 1.27 mm. The nightglow distribution is highly inhomogeneous with the regions of brightest emission generally located at low latitudes near the midnight meridian. Unexpectedly some orbits show also intense airglow activity over the south polar region. The spatially resolved airglow is spectacularly variable not only in its morphology and intensity but also in the apparent motions of the airglow small- and large-scale structures. Visual tracking of the bright features allowed to obtain mean zonal and meridional motions related to the subsolar to antisolar circulation. The zonal velocity is dominated by an intense prograde jet (contrary to the retrograde planetary rotation) from dawn to midnight extending up to 22 hours in local time with lower velocities and reversed sign from dusk. Typical zonal velocities range between +60 (prograde) to 50 (retrograde) m/s, whereas most meridional velocities range from 20 (poleward) to +100 m/s (equatorward) with an average meridional circulation of +20 m/s toward low latitudes. The brightest small-scale (100 km) features appear correlated with locations of apparent convergence which may be a signature of compression and downwelling, whereas this is not evident for the largescale structures suggesting slow subsidence over large areas mixed with horizontal motions. We argue that part of the tracked motions are representative of real motions at the mesosphere over an altitude range of 95–107 km. Citation: Hueso, R., A. Sa´nchez-Lavega, G. Piccioni, P. Drossart, J. C. Ge´rard, I. Khatuntsev, L. Zasova, and A. Migliorini (2008), Morphology and dynamics of Venus oxygen airglow from Venus Express/Visible and Infrared Thermal Imaging Spectrometer observations, J. Geophys. Res., 113, E00B02, doi:10.1029/2008JE003081.
1. Introduction [2] The circulation of the upper atmosphere of Venus is traditionally decomposed into two distinct flow patterns at different altitudes: A subsolar-to-antisolar (SS-AS) thermospheric circulation cell driven by solar heating, stable at levels above the mesosphere (z > 120 km [Bougher et al., 1997], and a superrotating zonal (SZ) flow in the planetary rotation sense at the upper troposphere (z 65 km) [see, e.g., Gierasch et al., 1997; Markiewitz et al., 2007; Limaye, 1 Departamento de Fı´sica Aplicada I, E.T.S. Ingenieros, Universidad del Paı´s Vasco, Bilbao, Spain. 2 Istituto di Astrofisica Spaziale e Fisica Cosmica, INAF-IASF Roma, Rome, Italy. 3 Laboratoire d’ Etudes Spatiales et d’ Instrumentation en Astrophysique, Observatoire de Paris, Universite´ Paris-Diderot, CNRS, UPMC, Meudon, France. 4 Laboratoire de Physique Atmosphe´rique et Plane´taire, Universite´ de Lie`ge, Lie`ge, Belgium. 5 Space Research Institute, Moscow, Russia.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003081$09.00
2007; Sa´nchez-Lavega et al., 2008]. In the transition region, the mesosphere (70 < z < 120 km), the interaction of both motions results in a complex circulation that greatly varies over time [Lellouch et al., 1997; Bougher et al., 2006]. The global circulation in the mesosphere is probably influenced by a number of physical processes such as breaking gravity waves [Alexander, 1992] and diffusive processes, but their effects over the mesospheric circulation are not well understood or quantified [Bougher et al., 2006]. Additionally, the return branch of the SS-AS upper circulation must lie somewhere in the mesosphere but its location has not been clearly identified. [3] While the dynamics of the mesosphere of Venus is hardly accessible to in situ measurements it can be studied by means of Doppler shifts in certain absorption lines sensitive to mesospheric levels [e.g., Clancy et al., 2007; Widemann et al., 2007] but this technique lacks spatial resolution, requires a sophisticated analysis of the observations and a precise treatment of systematic errors. A complimentary technique suited for exploration from a spacecraft is the study of nightglow molecular emissions which trace the motions of different chemical species in the
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upper atmosphere with the drawback that the emissions are coupled with photochemical processes. [4] Among the many nonthermal emissions produced in Venus, the oxygen airglow at 1.27 mm is the most intense. Its was first detected in ground-based observations of Venus by Connes et al. [1979]. This airglow is produced by the recombination of oxygen atoms dissociated by photolysis of CO2 at thermospheric altitudes on the sunlit hemisphere. In the upper mesosphere (95 – 110 km) three-body recombination of O atoms leads to O2 formation in a particular excited state, a1Dg, which is followed by airglow emission as the molecule relaxes to its X 3Sg ground state. The radiative lifetime of O2 singlet state is 70 min. [Lellouch et al., 1997; Miller et al., 2001]. The O2 airglow is observed in the night-side hemisphere at 1.267 microns [Allen et al., 1992; Bougher and Borucki, 1994] and with far less intensity in O2 Herzberg II visible wavelengths [Krasnopolsky et al., 1976; Bougher et al., 1997]. [5] It was soon realized that the intense oxygen IR nightglow is highly variable from day to day [Crisp et al., 1996], its intensity distribution is generally not symmetric in latitude [Allen et al., 1992], often exhibits multiple local maxima, and shows variations on timescales as short as 1 hour [Lellouch et al., 1997]. The related nitric oxide nightglow emission was mapped during the Pioneer Venus mission [Stewart et al., 1980] statistically presenting larger brightness close to the Equator and shifted by about 2 hours from midnight toward dawn. Both nightglows are decoupled in height providing information about different vertical levels: 95– 105 km for O2 airglow (G. Piccioni et al., Venus oxygen airglow vertical profile, submitted to Journal of Geophysical Research, 2008) and 115 km for NO [Ge´rard et al., 2008b]. [6] In this work we have used night side nadir observations from the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) [Drossart et al., 2007a] on board the Venus Express orbiter [Titov et al., 2006; Svedhem et al., 2007] to construct maps of the Venus atmosphere in the O2 emission band. The airglow maps were used to study the morphology and temporal variability of the airglow. In those orbits where the airglow is particularly intense we also measured the displacement of ‘‘cloud-like’’ airglow features. [7] Preliminary measurements of apparent motions were presented by Drossart et al. [2007b] for selected orbits. Ge´rard et al. [2008a] gave a precise account of the average distribution of the O2 nightglow. Other works in this special section (Piccioni et al., submitted manuscript, 2008) present results obtained in limb view which lead to vertical profiles of the airglow emission.
2. Observations and Methods [8] We analyzed data obtained during 28 consecutive orbits between 1 July 2006 and 28 July 2006, one orbit on 6 August 2006 presenting a good view of the equator, six orbits between December 2006 and January 2007 and five orbits from 7 April 2007 to 11 April 2007 that cover the polar region in higher detail than the rest of the VIRTIS data set. The data were selected according to the availability of high-resolution observations, fast repetition in each orbit (two or more images of the same region separated by 1hour) and the strong airglow emission present during most of the
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selected orbits. The Sun F10.7 flux (a convenient measurement of the solar activity in radio wavelengths commonly used as a proxy of the Sun’s irradiance in the Extreme UltraViolet) averaged over July 2006, August 2006, January 2007 and April 2007 was 76 10 22, 79 10 22, 84 10 22 and 72 10 22 W/m2 Hz respectively, reflecting homogeneous conditions for the Sun’s activity close to the solar minimum. [9] Oxygen airglow at 1.27 mm superimposes with thermal radiation coming from the lower atmosphere and partially filtered by the lower clouds. In order to subtract the thermal contribution from the total airglow signal we compare VIRTIS fluxes measured in the 1.26 – 1.28 mm wavelength range to the flux on the 1.18 mm observation window which provides a nearby region with flux dominated by the lower thermal emission. In addition, O2 photons emitted downward are subsequently backscattered by the underlying clouds and enhance the measured airglow radiation. A correction factor of 2.7 was used for nadir observations and geometry effects were corrected to first order by multiplying each pixel by the cosine of the emission angle [Crisp et al., 1996]. This crude estimation of brightness emission is adequate for VIRTIS observations of subpolar and midlatitudes (emission angles lower than 40°) but may fail in the few observations containing information about the tropics and the equator. Finally, depending on the latitudes being sampled we projected the images in longitude-latitude maps or polar stereographic projections to compare images obtained from different observational geometries. [10] Images separated by 1 hour allowed to track and retrieve the displacements of the airglow features in 23 of the 40 orbits, typically the orbits with higher intensity airglow structures, being the remaining orbits characterized by low intensity of the airglow or low contrast of the observed structures. The apparent motions of distinct airglow features were measured visually using the PLIA software [Peralta et al., 2005].
3. Morphology of the Oxygen Airglow and Temporal Variability 3.1. Low and Mid Latitudes [11] As a consequence of the quasi-polar VEX orbit, the airglow activity at low latitudes is only partially covered by VIRTIS observations with a resolution of 100 km pix 1 in some cases, while mid to subpolar latitudes are observed on many orbits with spatial resolution down to 15 km pix 1. [12] The statistical distribution of Venus oxygen nightglow is inhomogeneous with the regions of brightest emission located at low latitudes near and dawnward of the midnight meridian [Ge´rard et al., 2008b]. Maps of airglow brightness distribution of different regions are presented in Figure 1 using data from orbits that sample nearly simultaneously most of the Southern Hemisphere. On this data set the equatorial region at midnight always presents high values of brightness. Large structures with typical sizes of 3000 km extending from the equator to midlatitudes are often observed. In most cases the large-scale structures observed on Figure 1a are similar to the global structures seen in the global map of airglow brightness presented by Ge´rard et al. [2008a] and constructed by averaging several
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Figure 1. Airglow intensity map obtained during orbits (a) 87 (16 July 2006) and (b) 91 (20 July 2006). The wide coverage of the maps is close to half the night-side hemisphere and was obtained using data from three different VIRTIS data cubes for each map. The mean brightness of the sampled regions close to the equatorial midnight can change from 5 MR to 1 in a single day but the large values are more common. hundred maps from different dates. Remarkably, though largely expected, the mean airglow brightness in VIRTIS images (centered at the same latitude and local time and covering areas as big as 1.5 107 km2) changes dramatically from one orbit to other. Changes observed on Figure 1 are sometimes observed on timescales as short as 1 day over the VIRTIS data set. On the contrary, large-scale structures as seen on Figure 1a are stable on timescales of 2 hours or longer indicating sustained formation of excited oxygen molecules. 3.2. Polar Latitudes [13] Ge´rard et al. [2008a] give an average value of O2 airglow brightness at 70°S of 1.2 MegaRayleighs (MR). In contrast, intense airglow activity can also arise over polar latitudes in particular orbits. This was unexpected since the polar regions are far from the antisolar point and should contain lower densities of oxygen atoms. Figure 2 shows two images of a sequence of VIRTIS observations of the most intense airglow brightness over polar latitudes and viewed with an emission angle lower than 20° for all the sequence. A gaussian fit to the brightness temporal evolution of this structure following its motion gives a time decay of 3 hours with a peak brightness of 7 MR over a 100 km
elongated structure, moving toward lower latitudes with a meridional velocity of v = 50 ± 10 m/s. The brightness decay in time of this structure was about 2.5 times the radiative lifetime of the oxygen airglow indicating continuous production of excited molecules over this small region. [14] The period from 7 to 11 April 2007 was used for detailed research of the polar area finding intense airglow activity in all the VIRTIS images and sustained during all orbits. Figure 3 shows a sequence of observations where apparent motions of the airglow lead to a polar to low latitudes apparent motion with v 50 ± 5 m/s. Detailed motions in the whole movie during this orbit (composed of 12 images at time steps of 30 min to 1 hour) shows that the airglow evolves by a combination of apparent motions at small scale and activation and deactivation of airglow emission structures over extended regions.
4. Apparent Motions and Mesospheric Circulation [15] Figure 4 shows apparent motions tracked over airglow structures on different dates, latitudes and local times. The panels show that the morphology of the airglow is related with the apparent motions tracked. Large-scale
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morning hours (0600 0200 local time) and reversed in direction from 1800 to 2000 local time. The meridional motions are more intense at subpolar latitudes (70°S) with average velocities of v 20 m/s, decreasing toward lower latitudes and reaching a minimum value of v 5 m/s at 15°S and increasing again at equatorial latitudes. [18] In the orbits with high intensities of the airglow emission (notably on the equatorial case and the five polar orbits) the displacements of the airglow brightness patterns seem to take place at different altitude levels since the images show a mixture of structures moving with different speeds that visually seem to superimpose. In support of this interpretation, limb observations of airglow activity some-
Figure 2. Polar projections and time evolution of the brightest airglow structure seen on polar latitudes. Longitudes are given in terms of local times. Observations correspond to orbit 80, 9 July 2006, and the time difference between both observations is 3 hours. structures move slowly on one direction (generally perpendicular to the airglow brightness front) while smaller-scale structures move more rapidly in the direction defined by the front (see Figure 4a). [16] The equatorial region covered on Figure 4d looks turbulent with bright features colliding suggesting that the observed motions may come from a variety of different altitude levels. Some of the brightest structures (notably on Figure 4e, but also on Figures 4c and 4d and particular regions on Figure 4a) correlate with locations where convergence of the apparent motions is present. This behavior may not apply to the larger-scale structures, since they appear surrounded by far less intense airglow features and thus, it is not possible to retrieve motions around them. [17] Figure 5 shows the mean horizontal circulation in Venus’ night-side hemisphere obtained by spatially and temporally averaging the ensemble of apparent motions obtained in 23 orbits with 1000 tracked motions. The upper part of Figure 5 also shows the local time-latitude location of the specific regions displayed in Figure 4. The large variability in intensity and direction of the velocities, which sometimes can be as high as 100 m/s, results in low values of the average speeds. The most remarkable feature in this plot is a repeatedly observed jet at 35°S with an average zonal speed of u = 30 m/s opposite to the planetary rotation and with motions twice as intense in the early
Figure 3. Polar projections and time evolution of the 1.267 mm radiance over orbit 355 obtained on 27 August 2007 at local times: (a) 1539, (b) 1739, (c) 1939, and (d) 2209. The brightest airglow features have an intensity of 1.8 MR on the first image and longitudes are given in terms of local times. The global apparent motions are toward higher latitudes. Some of the structures like the bright arm at 68°S in Figure 3d appear suddenly and are not related with horizontal motions but with the sudden airglow brightening of new regions.
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Figure 4. Average apparent motions and correlation with the airglow intensity for selected orbits: (a) orbit 96 obtained on 25 July 2006; (b) orbit 99 obtained on 28 July 2006; (c) orbit 76 obtained on 5 July 2006; (d) orbit 108 obtained on 6 August 2006; (e) orbit 84 obtained on 13 July 2006; and (f) orbit 80 obtained on 9 July 2006. The maps cover a significant portion of the local time-latitude plane showing repeatedly observed motions, like the zonal jet at 30°S – 45°S opposite to the planetary rotation (Figures 4a, 4b, 4e, and 4f). Opposite motions between 45°S and 75°S that seem part of a close circulation are observed at different local times on different orbits (Figure 4b); meridional motions are toward the pole (Figure 4a) but generally directed toward low latitudes (Figures 4a, 4b, 4c, and 4f). The few observations of the equator show chaotic motions with a mean prograde (in the planetary rotation sense) component (Figure 4d). Brightness emission given are only approximative for Figure 4d with emission angles 70°. times presents peaks at two different altitude levels, one at 96 ± 1 km altitude and a second peak at 103 –107 km altitude (Piccioni et al., submitted manuscript, 2008).
5. Discussion [19] The average motions obtained are in agreement with a subsolar to antisolar circulation at the IR oxygen airglow levels (96 km) with evident meridional motions from polar to low latitudes and large variability of the zonal and meridional components. [20] Caution must be taken in interpreting the airglow brightness patterns displacements as tracers of flow motions in the lower mesosphere. They could be influenced by the activation and/or deactivation of airglow emission at specific regions by other phenomena (i.e., gravity waves affecting atmospheric density, localized or extended regions of downwelling, atmospheric diffusivity, etc). [21] An additional level of complexity is introduced by the radiative lifetime (70 min) of the O2 singlet state which introduces difficulties in the interpretation of the observed rates of brightness variation. The e-folding time constant of any local change (decrease or increase) cannot be less than 70 min without advection of air masses containing different proportion of glowing O2 molecules. For airglow brightness decreasing more slowly the brightness reduction could be
caused by a combination of horizontal transport and a decrease of the flux of O atoms modifying the local O density and airglow brightness. In those locations where the airglow intensity increases there must be enhanced downward flux of O atoms. [22] The observations show two types of behavior for the Venus oxygen airglow temporal evolution. On the one hand, the large-scale 1000 –3000 km structures survive at least for a few hours (longer than the radiative decay time) but change dramatically from day to day. Since limb scans of O2 emission slightly depend on latitude with higher latitudes characterized by higher altitudes of the airglow (Piccioni et al., submitted manuscript, 2008) these large structures may descend progressively as they travel from polar latitudes to the equator, compressing and increasing the concentration of free O atoms. On the other hand, smallscale bright structures with smaller scales (100 km) seem related with the apparent convergence of motions tracked in the airglow. In fact, the time evolution of the airglow maps seems composed of two types of motions, the large-scale structures moving like ‘‘fronts’’ or waves, and the bright structures immersed on them or isolated, apparently moving like passive tracers. We argue that at least part of the later motions are real motions at the mesosphere as the newly formed oxygen molecules are advected by the mesospheric circulation.
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Figure 5. Mean apparent circulation of the mesospheric oxygen airglow brightness. (a) Averages of all apparent motions tracked on the airglow and also marks the sampled regions appearing on Figure 4. A reference vector of 100 m/s is included on the upper left corner of the map. (b) A schematic model of the mean circulation of Venus mesosphere from these measurements. The background image corresponds to a map of airglow brightness obtained on orbit 93, 22 July 2007, that can be compared with those on Figure 1. Regions with no arrows indicate regions with not much available data. The apparent rotating motions indicated at 50°S are found at a variety of local times depending on the particular orbit. The jet from dawn to midnight is a repetitive structure while the jet from dusk to midnight is less regular from one orbit to another. [23] The brightest airglow patches may be regions of strong subsidence (and local adiabatic warming) associated with downdrafts that increase the volume concentration of O atoms. If this is correct, spectral maps of the O2 IR nightglow should show a positive correlation between the brightness and the rotational temperature of the emission line. Partial confirmation of this point comes from stellar occultation experiments performed by Venus Express [Bertaux et al., 2007] which result on warm temperatures at the airglow emission levels (95 – 100 km). Ground-based observations have been made to correlate the rotational temperature and brightness with results sometimes indicating such a correlation [Ohtsuki et al., 2005] and sometimes not with some dependence on the morphology of the airglow region [Bailey et al., 2007]. [24] Although the present work does not fully disentangle the different contributions to the airglow variability and therefore is not able to fully retrieve the mesospheric true motions it does show that such an approach is possible.
VIRTIS-M observations obtained with shorter time differences than the radiative lifetime of the oxygen airglow (70 min) will partially help to solve this problem. Although part of the airglow motions tracked on this work were obtained on image pairs separated by 30 min the short time difference between the images translate into larger errors in the derived motions and most of the data corresponds to images separated by 1 hour. Both data sets produce similar results on terms of the mean derived circulation. [25] Further analysis should concentrate on identifying any correlation between local temperature anomalies and O2 intensities. Also, detailed comparison with Venus GCMs should allow to better constrain the role of gravity waves as drivers of the airglow variability. The Venus Express Extended mission should run during 2008 – 2009 with higher levels of solar activity. According to NOAA Space Weather Prediction Center (SWPC) a reasonable estimate is an increase of a 60% of the F10.7 index for the end of 2009 (http://www.sec.noaa.gov/SolarCycle/SC24). This
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will allow a higher airglow activity and a better sampling of Venus mesosphere through its spatial distribution and variability. [26] Over the next years the combination of results from Venus Express instruments (VIRTIS O2 IR maps, SPICAV thermal profiles and NO nightglow observations, VeRa thermal data combined with cyclostrophic balance over mid to high latitudes and cloud top wind tracking with VMC and VIRTIS) and ground-based observations (nightglows and Doppler winds retrieval) will allow a significant improvement of our understanding of the mesospheric dynamics and the underlying mechanisms driving its variability. [27] Acknowledgments. Venus Express is a mission of the European Space Agency. We gratefully thank all members of the ESA Venus Express project and of the VIRTIS technical team. This work has been funded by Spanish MEC AYA2003-03216, with FEDER support and Grupos UPV 15946/2004 and is supported by the CNES and ASI national space agencies. R.H. acknowledges a Ramo´n y Cajal contract from MEC. J.C.G. acknowledges funding from the Belgian Fund for Scientific Research and the PRODEX program managed by the European Space Agency with the help of the Belgian Federal Space Science Office.
References Alexander, M. J. (1992), A mechanism for the Venus thermospheric superrotation, Geophys. Res. Lett., 19, 2207 – 2210. Allen, D., D. Crisp, and V. S. Meadows (1992), Variable oxygen airglow on Venus as a probe of atmospheric dynamics, Nature, 359, 516 – 519. Bailey, J. A., V. S. Meadows, S. Chamberlain, A. Simpson, and D. Crisp (2007), Variability of the Venus Oxygen Airglow, Bull. Am. Astron. Soc., 38, 526. Bertaux, J. L., et al. (2007), A warm layer in Venus’ cryosphere and highaltitude measurements of HF, HCl, H2O and HDO, Nature, 450, 646 – 649. Bougher, S. W., and W. J. Borucki (1994), Venus O2 visible and IR nightglow: Implications for lower thermosphere dynamics and chemistry, J. Geophys. Res., 99, 3759 – 3776. Bougher, S. W., M. J. Alexander, and H. G. Mayr (1997), Upper atmosphere dynamics: Global circulation and gravity waves, in Venus II: Geology, Geophysics, Atmospheres, and Solar Wind Environment, edited by S. W. Bougher, D. M. Hunten, and R. J. Philips, Univ. of Ariz. Press, Tucson. Bougher, S. W., S. Rafkin, and P. Drossart (2006), Dynamics of the Venus upper atmosphere: Outstanding problems and new constraints expected from Venus Express, Planet. Space Sci., 54, 1371 – 1380. Clancy, R. T., B. J. Sandor, and G. H. Moriarty-Schieven (2007), Dynamics of the Venus upper atmosphere: Global-temporal distribution of winds, temperature, and CO at the Venus mesopause, Bull. Am. Astron. Soc., 39, abstract 61.07. Connes, P., J. F. Noxon, W. A. Traub, and N. P. Carleton (1979), O2 (1D) emisssion in the day and night airglow of Venus, Astrophys. J., 233, L29 – L32. Crisp, D., V. S. Meadows, B. Be´zard, C. deBergh, J.-P. Maillard, and F. P. Mills (1996), Ground-based near-infrared observations of the Venus nightside: 1.27 mm O2 (a1Dg) airglow from the upper atmosphere, J. Geophys. Res., 101, 4577 – 4594. Drossart, P., et al. (2007a), Scientific goals for the observation of Venus by VIRTIS on ESA/Venus express mission, Planet. Space Sci., 55, 1653 – 1672. Drossart, P., et al. (2007b), Venus upper atmospheric emissions from VIRTIS spectral imaging observations, Nature, 450, 641 – 645.
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Ge´rard, J.-C., A. Saglam, G. Piccioni, P. Drossart, C. Cox, S. Erard, R. Hueso, and A. Sa´nchez-Lavega (2008a), Distribution of the O2 infrared nightglow observed with VIRTIS on board Venus Express, Geophys. Res. Lett., 35, L02207, doi:10.1029/2007GL032021. Ge´rard, J.-C., C. Cox, A. Saglam, J.-L. Bertaux, E. Villard, and C. Nehme´ (2008b), Limb observations of the ultraviolet nitric oxide nightglow with SPICAV on board Venus Express, J. Geophys. Res., doi:10.1029/ 2008JE003078, in press. Gierasch, P. J., et al. (1997), The general circulation of the Venus atmosphere: An assessment, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher, D. M. Hunten, and R. J. Philips, pp. 459 – 500, Univ. of Ariz. Press, Tucson. Krasnopolsky, V. A., A. A. Krysko, V. N. Rogachev, and V. A. Parshev (1976), Spectroscopy of the Venus night airglow from the Venera 9 and 10 satellites, Cosmic Res., 14, 789 – 795. Lellouch, E., T. Clancy, D. Crisp, A. Kliore, D. Titov, and S. W. Bougher (1997), Monitoring of mesospheric structure and dynamics, in Venus II: Geology, Geophysics, Atmospheres, and Solar Wind Environment, edited by S. W. Bougher, D. M. Hunten, and R. J. Philips, pp. 295 – 324, Univ. of Ariz. Press, Tucson. Limaye, S. S. (2007), Venus atmospheric circulation: Known and unknown, J. Geophys. Res., 112, E04S09, doi:10.1029/2006JE002814. Markiewitz, W. J., D. V. Titov, S. S. Limaye, H. U. Keller, N. Ignatiev, R. Jaumann, N. Thomas, H. Michalik, R. Moissl, and P. Russo (2007), Morphology and dynamics of the upper cloud layer of Venus, Nature, 450, 633 – 635. Miller, H. C., J. E. McCord, J. Choy, and G. D. Hager (2001), Measurement of the radiative lifetime of O2 (a1Dg) using cavity ring down spectroscopy, J. Quant. Spectrosc. Radiat. Transfer, 69, 305 – 325. Ohtsuki, S., N. Iwagami, H. Sagawa, Y. Kasaba, M. Ueno, and T. Imamura (2005), Ground-based observation of the Venus 1.27-mm O2 airglow, Adv. Space Res., 36, 2038 – 2042. Peralta, J., R. Hueso, N. Barrado, and A. Sa´nchez-Lavega (2005), Introducing PLIA: The planetary laboratory for image analysis, Bull Am. Astron. Soc., 37, 653. Sa´nchez-Lavega, A., et al. (2008), Variable winds on Venus mapped in three dimensions, Geophys. Res. Lett., 35, L13204, doi:10.1029/ 2008GL033817. Stewart, A. I. F., J. C. Gerard, D. W. Rusch, and S. W. Bougher (1980), Morphology of the Venus ultraviolet night airglow, J. Geophys. Res., 85, 7861 – 7870. Svedhem, H., et al. (2007), Venus Express—The first European mission to Venus, Planet. Space Sci., 55, 1636 – 1652. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297. Widemann, T., E. Lellouch, and A. Campargue (2007), New wind measurements in Venus’ lower mesosphere from visible spectroscopy, Planet. Space Sci., 55, 1741 – 1756. P. Drossart, Laboratoire d’ Etudes Spatiales et d’ Instrumentation en Astrophysique, Observatoire de Paris, Universite´ Paris-Diderot, CNRS, UPMC, 5 place Jules Janssen, F-92195 Meudon, France. (pierre.drossart@ obspm.fr) J. C. Ge´rard, Laboratoire de Physique Atmosphe´rique et Plane´taire, Universite´ de Lie`ge, 5 Avenue de Cointe, B-4000 Lie`ge, Belgium. (
[email protected]) R. Hueso and A. Sa´nchez-Lavega, Departamento de Fı´sica Aplicada I, E.T.S. Ingenieros, Universidad del Paı´s Vasco, Alameda Urquijo s/n, E-48013 Bilbao, Spain. (
[email protected];
[email protected]) I. Khatuntsev and L. Zasova, Space Research Institute, Profsouznaya 84/ 32, Moscow, Russia. (
[email protected];
[email protected]) A. Migliorini and G. Piccioni, Istituto di Astrofisica Spaziale e Fisica Cosmica, INAF-IASF Roma, Via del Foso del Cavaliere 100, I-00133 Rome, Italy. (
[email protected]; giuseppe.piccioni@ iasf-roma.inaf.it)
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Limb observations of CO2 and CO non-LTE emissions in the Venus atmosphere by VIRTIS/Venus Express G. Gilli,1 M. A. Lo´pez-Valverde,1 P. Drossart,2 G. Piccioni,3 S. Erard,2 and A. Cardesı´n Moinelo3 Received 11 February 2008; revised 29 October 2008; accepted 13 November 2008; published 7 March 2009.
[1] We report and analyze here observations of strong infrared emissions from the limb of
the Venus upper atmosphere during daytime, taken by the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) aboard Venus Express. We focus on the measurements taken during the first 4 months of nominal operations. The emissions observed at 4.3 mm and at 2.7 mm are attributed to CO2 fluorescence of solar radiation and are detected up to about 160 km and 130 km, respectively, while the CO fluorescence at 4.7 mm is observed up to about 120 km. The emissions are detected in both the channels of VIRTIS, at different spatial and spectral resolutions (resolving powers about 1800 and 400), for the periapsis and the apoapsis of the Venus Express orbit. From these data sets we built up 2-D maps of the emissions as well as vertical profiles, which are then studied in order to characterize their variations with geophysical parameters, like solar illumination and emission altitude. Several analyses are performed in order to understand the VIRTIS behavior, to determine systematic effects in the data, and to propose appropriate corrections. We also present comparisons with a theoretical nonlocal thermodynamic equilibrium (non-LTE) model of the Venus upper atmosphere. The agreement is very encouraging, in general, and the main variability observed in the data, with solar zenith angle and altitude, can be understood with the model. We conclude that the present data set opens brilliant perspectives for deriving densities and rotational temperatures in the upper mesosphere and lower thermosphere of Venus. Citation: Gilli, G., M. A. Lo´pez-Valverde, P. Drossart, G. Piccioni, S. Erard, and A. Cardesı´n Moinelo (2009), Limb observations of CO2 and CO non-LTE emissions in the Venus atmosphere by VIRTIS/Venus Express, J. Geophys. Res., 114, E00B29, doi:10.1029/2008JE003112.
1. Introduction [2] The European mission Venus Express arrived at Venus in April 2006, and started its scientific phase in June 2006, being the first mission in more than 20 years to study the Venus atmosphere systematically from orbit [Svedhem et al., 2007]. In contrast to the Earth’s mesosphere and lower thermosphere, the equivalent density layers on Venus from about 90 to 150 km, are regions scarcely observed using remote sounding in the infrared (IR). The most abundant species CO2, is well known to present strong IR vibrationalrotational bands, and more than 25 years ago, Pioneer Venus made systematic nadir observations using the fundamental band of CO2 at 15 mm in order to retrieve atmospheric temperatures in the mesosphere of Venus [Taylor et al., 1980]. Also, NIMS/Galileo permitted a sounding of the lower mesosphere temperature at 4.3 mm in nadir and nighttime [Roos-Serote et al., 1995]. However, at higher altitudes, these emissions are expected to be out of thermo1 2 3
Instituto de Astrofı´sica de Andalucı´a, CSIC, Granada, Spain. Observatoire de Paris, Meudon, France. IASF, INAF, Rome, Italy.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003112$09.00
dynamic equilibrium, with particularly large deviations during daytime, owing to solar pumping of the corresponding vibrational states [Dickinson, 1972; Deming et al., 1983; Rolda´n et al., 2000]. As a consequence, their inversion in terms of geophysical parameters is not straightforward, or in other words, their thermodynamic information content is limited [Lo´pez-Puertas and Taylor, 2001]. The understanding of these non-LTE emissions is important, not only for its own interest, but also they can give rise to a large radiative heating, or cooling, of the atmosphere. In the case of Venus, they are a key ingredient of the energy balance between 80 and 150 km [Bougher et al., 1994]. Rolda´n et al. [2000] showed that the CO2 bands at 4.3 mm and the ones at 2.7 mm dominate the solar heating at those altitudes; much less important are the CO absorptions at 4.7 mm, given the lower atmospheric abundance of CO. Further observational problems challenged an infrared sounding of these layers from the Venus orbit so far; among them, the pointing in the limb (the optimum geometry of observation at these low atmospheric densities) and the sensitivity of the infrared detectors. [3] Progress toward limb sounding of planetary atmospheres has been made during the last 2 decades, both theoretically and observationally. Triggered by ground-based
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observations of CO2 laser bands at 10 mm, a number of nonLTE models for atmospheric IR emissions were developed for Mars and Venus in the 1980s [Deming et al., 1983; Gordiets and Panchenko, 1983; Stepanova and Shved, 1985]. Later, more comprehensive tools were developed in preparation for sounding of the Martian atmosphere in the IR during the 1990s [Lo´pez-Valverde and Lo´pez-Puertas, 1994], which were adapted later to the Venus atmosphere [Rolda´n et al., 2000]. These models, which included many more radiative bands and CO2 energy states, represent adequate tools for simulation and analysis of limb data. Some of the few non-LTE observations available so far, and amenable for such studies, were taken by the Near Infrared Mapping Spectrometer (NIMS) instrument [Carlson et al., 1992] on board the Galileo sounder, during its flyby of Venus in 1990 [Carlson et al., 1991]. A few limb spectra of the CO2 4.3-mm band were obtained during daytime and were explained recently using an improved version of the non-LTE model of Rolda´n and colleagues [Lo´pez-Valverde et al., 2007] (hereinafter LVEA). Recent analysis of CO2 non-LTE emissions have also been carried out for Mars using data from the Planetary Fourier Spectrometer (PFS) on board Mars Express [see Lo´pez-Valverde et al., 2005; Formisano et al., 2006], which helped to partially validate the non-LTE model. Regarding CO, non-LTE emission lines in the near-IR were first discovered in the Venus atmosphere about 20 years ago by de Bergh et al. [1988]. They were recently modeled and analyzed by Crovisier et al. [2006], who identified the fundamental CO(1-0) and the first hot CO(2-1) vibrational-rotational bands, derived rotational temperatures from them, and predicted that the instrument VIRTIS/Venus Express might detect these emissions. [4] VIRTIS is one of the instruments aboard Venus Express with capabilities to sound the upper atmosphere of Venus [European Space Agency (ESA), 2001; Titov et al., 2006]. It is an imaging spectrometer in the visible and near infrared, inherited partly from the Rosetta mission [Coradini et al., 1998]. VIRTIS is devoted to a large number of investigations at Venus, like the cloud deck morphology and dynamics, the lower atmospheric composition, and the derivation of the temperature structure in the lower mesosphere. First studies of these aspects are presented in companion papers in this issue. Most of these studies are carried out by nadir observations. However, VIRTIS is designed to perform limb observations aboard Venus Express as well [Titov et al., 2006]. This mode of observation represents the first systematic sounding of the upper layers of the Venus atmosphere in the infrared, and is devoted to understand the Venus high-atmosphere emissions [Drossart et al., 2007b]. The detection of the non-LTE CO2 4.3-mm band by VIRTIS at thermospheric altitudes has been confirmed by Drossart et al. [2007a], in consonance with the theoretical expectations. [5] In this paper, we focus on a detailed and systematic analysis of VIRTIS observations of CO2 and CO non-LTE emissions from the upper atmosphere of Venus using such a limb geometry. This is a unique data set, and our goals include, first of all, performing an internal validation of the data of our interest, second, characterizing the behavior of the instrument during limb sounding, and also describing the data set available, its quality and quantity, and its
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scientific potential. These CO2 non-LTE emissions have also been observed by VIRTIS in nadir, which show signatures of gravity waves propagation into the thermosphere, as studied in a companion work in this issue [Garcı´a et al., 2009]. The study of the O2 infrared system in the Venus mesosphere is also presented by Hueso et al. [2008] and in this issue (G. Piccioni et al., Near IR oxygen nightglow observed by VIRTIS in the Venus upper atmosphere, submitted to Journal of Geophysical Research, 2009). [6] In this paper we also present and discuss comparisons with the non-LTE model results in LVEA, using both their published results and a small number of specific simulations performed here with such model. Although a more detailed comparison between model and data is under preparation, this is the first time that this model is confronted with an extensive data set of such a high spectral and spatial resolution. This work is, therefore, also devoted to test the model’s main theoretical predictions, and to identify improvements and lines of exploration for non-LTE models. [7] In section 2, we first describe VIRTIS observations in the limb, including the presentation of 2-D maps and vertical profiles of radiances. Their validation analysis, including consistency tests, actual noise level determination, as well as internal correlations and comparisons with the non-LTE model, are presented in section 3. Further discussions, future prospects and conclusions, are presented in section 4.
2. VIRTIS Observations [8] VIRTIS on Venus Express is an imaging spectrometer whose precursor is currently en route to comet 67P/ Churyumov-Gerasimenko, as part of the scientific payload of the Rosetta mission. Its detailed description and calibration can be found elsewhere [Coradini et al., 1998; Piccioni et al., 2006; Titov et al., 2006] and see also the companion papers in this issue. We briefly describe here those characteristics and mode of operation on board Venus Express which are needed for our analysis. 2.1. VIRTIS/VEX Characteristics and Data Set [9] The instrument consists of two channels: VIRTIS-M (hereinafter V-M), a mapping spectrometer at a resolving power R 400, working in the visible (0.3 –1 mm) and in the near infrared (1 – 5 mm), and VIRTIS-H (hereinafter V-H), a high-resolution spectrometer (R 1800) working in the spectral range 2 –5 mm. Table 1 of the work by Drossart et al. [2007b], summarizes their main characteristics; let us recall some of them here. The nominal field of view (FOV) is 64 64 mrad2 for V-M (whole frame, with 256 spatial pixels in each direction), and 0.58 1.75 mrad2 for V-H. The V-M spectral range is sampled at 432 spectral wavelengths, while a typical V-H spectrum obtained in the nominal mode contains a sequence of 3456 measurements. The nominal noise equivalent radiance (NESR), for 1 second of integration time, is 5000 mW m 2 sr 1 mm 1 in both channels, according to the ESA [2001]. All V-H and V-M data taken during one orbit, or terrestrial day, are stored in the so-called ‘‘qubes.’’ The V-M ones generally contain a set of data in 3 dimensions, 2 spatial and 1 spectral. During each orbit the number of V-H/V-M data qubes
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Figure 1. Geographical coverage with VIRTIS-M (V-M) during orbit 25, from apoapsis. (top, left) Map of radiances at 4.32 mm from qubes 0, 1, and 2. (top, right) Altitude-SZA location of all the limb data above the cloud layer (all pixels of seven qubes of this orbit); (bottom) Location of the data in SZA, longitude, latitude, and local time. The numbers identify the different qubes. obtained varies from 1 to 20. A detailed description of data handling and archiving is presented by Drossart et al. [2007b]. [10] VIRTIS observations are separated into two categories, corresponding to distances of Venus Express to the planet lower than 12,000 km (Spectral Mode) or higher (Spectral Imaging Mode). During the Spectral Mode only a partial coverage of the surface is normally obtained, but thanks to the proximity to the planet, a very small field of view (projected) is achieved, ideal to study atmospheric variability and composition, and for tangential limb observations. Several instruments on board Venus Express share this operation mode. The Spectral Imaging Mode permits unique and unprecedented maps of the disk of Venus, and has been used to obtain global observations from apoapsis (‘‘mosaic construction’’). We will show below how these VIRTIS 2-D images contain useful information on the limb of the planet as well. [11] A note regarding the distinction between limb and nadir observations is needed in order to clarify that, in a general sense, a data qube from VIRTIS will contain both kind of observations. Our distinction merely tells whether the line of sight touches the disk of Venus (nadir data) or crosses its atmospheric limb. Therefore, some pixels may represent actual limb sounding while nearby pixels in the same V-M image will supply ‘‘disk’’ observations. This will become clearer in section 2.5. For this reason, analysis of nadir observations, described in companion papers of this issue, share similar difficulties to our study and all of them are beneficial to the overall validation of VIRTIS data.
2.2. Geographical and Temporal Coverage [12] At present, V-M data are available in calibrated form up to the orbit 631 (13 May 2006 to 1 Jan 2008), but in this work we focused our analysis on a selected number of orbits up to orbit 295 (10 February 2007). For V-H, at the time of preparing this paper, the data from orbit 23 to 127 (13 May 2006 to 25 August 2006) were already calibrated, so we focused on them, while only a fraction was calibrated and available for the rest of the mission. As mentioned above, our purpose is to perform an exploration of V-M and V-H limb data in the Venus upper atmosphere, from the cloud top, around 60 km, up to upper thermospheric layers. With this aim, we made a catalog of orbits/qubes/pixels with actual limb sounding and a number of diagrams of the geometrical and spatial coverage of all limb data obtained. Examples are shown and commented next. 2.2.1. VIRTIS-M [13] Starting with V-M, Figures 1 and 2 show two examples of the V-M coverage for two different observational geometries, both sampling the Venus limb, one from orbit 25 (15 May 2006), during the apoapsis, and the other from orbit 43 (2 June 2006) near the periapsis. They can be considered as representative of the two principal VIRTIS modes of observation (see section 2.1). [14] The first orbit, from apoapsis, contains 8 qubes of V-M data (7 with limb data), each one consisting of set of pixels, or a 2-D image at each wavelength, as mentioned above. Let us recall that each pixel is associated to one individual spectrum. In Figure 1a, data at the wavelength of 4.32 mm from three of the qubes, building up a map of radiances; we will discuss them below. Each qube includes a
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Figure 2. V-M coverage in tangent altitude versus SZA during periapsis of orbits 43 and 287. Data are shown between 60 and 180 km only. large number of pixels (245 256), and the fraction of those pixels with actual pointing to the Venus limb are shown in Figures 1b– 1d. We see that when the 8 qubes are considered, the latitude, longitude, and local time of the data cover an ample range of values. In this work we pay particular attention to the altitude and solar zenith angle (SZA) coverage. These two parameters are expected to produce the maximum variation in the non-LTE emission of both CO2 and CO in the atmosphere, according to the model of LVEA. The most intense signal is expected between 90 and 140 km, and for low SZA values; for nighttime the signal would be very low, and we confirmed all of this with VIRTIS data. Figure 1b shows that the number of limb daytime spectra is large for this orbit. The vertical resolution from each row of pixels (for each qube) is between 15 and 20 km, typically. This is larger than an atmospheric scale height, but the superposition of adjacent rows may improve significantly this resolution, as we will show below. Also the SZA sampling is not completely even, but the large density of points permits suitable averaging if they are needed. [15] Figure 2 illustrates the V-M coverage in the qube 0 of orbit 43 (2 June 2006), during the periapsis of Venus Express and in the qubes 0 and 1 of the orbit 287 (1 February 2007). We use again altitudes at the tangent point versus SZA, which describe our preferred data set for limb sounding. The available daytime limb data during the orbit 43 are about 5410 while this number is larger for the 287 (about 8000). The vertical resolution obtained at the limb from one single row of pixels (300 m), improves by a very large factor with respect to the apoapsis case, and is therefore particularly useful for studying the vertical variation of the non-LTE emissions and testing the theoretical predictions. 2.2.2. VIRTIS-H [16] Observations of the limb of the planet by VIRTIS in the Spectral Mode are not as frequent as in the Spectral
Imaging Mode owing to pointing constraints; the field of view of the V-M frame is much larger than the V-H. The amount of V-H spectra which actually touches the limb of the planet is therefore very limited in most orbits, sometimes only 2 or 3 spectra of our interest are available per orbit. [17] Figure 3 illustrates the V-H geometrical coverage at the limb, again as a function of SZA and tangent altitude, combining all the orbits during the first four months of operation. A total of 651 points are shown, each one representing a V-H spectrum. We have marked 4 specific orbits (43, 76, 80, and 88) to indicate 4 different special observational coverages. The first one corresponds to a special limb case, the so-called ‘‘Tangential Limb Scenario’’ [see Titov et al., 2006], where the satellite is fixed pointing to the limb of Venus during the periapsis. This is a specially useful kind of orbit for limb sounding and for non-LTE studies, given the high rate of tangent point observations. Orbit 76 (6 July 2006) corresponds to a ‘‘stellar occultation’’ case, where again a large number of observations is obtained at the limb, but they correspond to the nighttime hemisphere. They do not show a distinct emission level and have not been considered in the analysis. The other two orbits correspond to the most common limb scenario. [18] For a number of studies, averaging of these data may be required, in order to achieve a good SZA or altitude coverage. From the 651 V-H limb spectra available, a total of 145 are daytime observations (SZA < 90°), and 44 of these correspond to strong solar illumination conditions (SZA < 60°). The last ones are, in principle, optimum for non-LTE studies. Let us recall that, within Venus Express, V-H data are the measurements with the highest spectral resolution. The latitude and longitude coverage of the daytime spectra are shown in Figure 3 as well. Except for orbits 43 (2 June 2006) and 80 (10 July 2006), with highlatitude sampling of nearby points, all others are disperse, with a preference for low latitude.
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Figure 3. (left) Coverage in tangent altitude and SZA of the VIRTIS (V-H) data available from all the orbits from 23 to 127. Some orbits are indicated with special symbols. (right) Latitude-longitude location of the data, for daytime only. Symbols as in the left. See section 2.2.2 for details. 2.3. VIRTIS-H Spectra and Profiles [19] We describe here typical V-H limb spectra which show strong non-LTE signals, and we build up vertical profiles also, in order to characterize the emissions. We focus on three emission bands, those at 4.3 mm and 2.7 mm both due to CO2, and that at 4.7 mm, due to CO. All of them are suitable to be contrasted with the recent non-LTE model predictions of LVEA. 2.3.1. CO2 4.3 mm Emission [20] Figure 4 (left) shows a sample of seven V-H spectra in the 4.3-mm region, corresponding to different orbits, and pointing at different tangent altitudes. The selection corresponds to orbits 33, 40, and 47 (24 May, 31 May, and 7 June 2006, respectively) and to a SZA interval from 20° to 30°. First, it is noticeable an intensity of the signal well above the noise level at all wavelengths between 4.20 and 4.50 mm. The emission falls to zero shortward and longward of these values, respectively, where the small oscillation observed can be considered as the measurement noise and which seems to be close to 2000 mW m 2 sr 1 mm 1. Only the two lowest altitude spectra show signal in the wings, though surely from solar scattering in the clouds. At the uppermost altitude shown in Figure 4, 167 km, the spectrum is still above noise level, which demonstrate that V-H can sound up to those thermospheric altitudes in Venus before averaging.
[21] It is clear that the spectral shape of the 4.3-mm band, the peak intensity, and the wavelength of the peak of the band, vary with altitude. The maximum emission is obtained in this sample around 110 km tangent altitude. This is close to the altitude predicted by the CO2 non-LTE model by LVEA. Also the spectral shape is similar to the model prediction, with three clear peaks in the 110 km spectrum, around 4.28, 4.32 and 4.35 mm. According to the model, the first peak is due to the second hot band of the main CO2 isotope; the mechanism being solar absorption at 2.7 mm, exciting CO2 states which later relax radiatively, emitting in the 4.3-mm region. The second peak, around 4.32 mm, is produced by radiative relaxation of the CO2 vibrational states excited after solar absorption at 2.7 mm and 2.0 mm. The third peak is mostly due to solar pumping at 2.0 mm. The non-LTE model predicts that at high tangent altitudes the 4.3-mm band would change this shape to a simpler two emission peaks separated by a dent around 4.29 mm, due to the dominance of the 2.7 mm solar pumping. This is indeed observed in Figure 4, above about 130 km. At higher altitudes, the model predicts a further change, due to the direct solar absorption in the fundamental band of the main isotope, whose central dip at 4.26 mm is observed in Figure 4 in the 167 km spectrum. [22] Vertical profiles can be constructed combining spectra at different altitudes. They, however, will not represent
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Figure 4. (left) Selection of seven V-H spectra from orbits 33, 40, and 47 at different altitudes and for 20° < SZA < 30°. (right) Vertical distribution of the 145 daytime measurements at 4.32 mm from orbits 23 to 127 (May – August 2006). The solid line on the right represents a model simulation. Orbits and symbols as in Figure 3. real 1-D profiles of the Venus atmosphere. Only a couple of orbits (43 and 80) contain a sufficient number of limb spectra to build a vertical profile for one single orbit; in general, data from different dates and geographical locations will be combined. This will have to be taken into account when performing detailed comparisons with non-LTE models, which usually provide the radiance structure at a single geographical location. In Figure 4 (right) we combine diverse data at 4.32 mm. The data and the symbols correspond to those in Figure 3. A tentative grouping by SZA values did not show any clear trend, given the scarcity of data; the SZA variation is better observed in V-M data and discussed below. [23] The cloud of V-H data presents a dispersion not present in the 1-D radiance simulations by LVEA, also shown in Figure 4. The model simulation corresponds to SZA = 80° similar to that of orbit 80. This orbit can be considered as a good 1-D profile extending from 30 to 200 km. Notice that data from orbit 43 can also be used to build vertical profiles, but only above about 100 km. Both data and model present a clear increase with altitude in the lower mesosphere and a peak emission around 120 km. Above this altitude, the emission declines quickly to zero, following the decrease in atmospheric density. The actual atmospheric density profile is critical in determining the precise altitude of the peak emission, and this is the most likely reason for the mismatch with the prediction of LVEA. Regarding the peak radiances, not much comparison can be performed at this stage, as it should be the topic of a more rigorous comparison of the non-LTE model, and should account for suitable atmospheric variability in the density profiles.
2.3.2. CO2 2.7 mm Emission [24] According to LVEA, CO2 emissions at 4.3 mm and at 2.7 mm are both produced by solar pumping. The emission level at 2.7 mm is, however, much lower and we have only found a small number of cases/spectra with a distinct emission above noise levels. Figure 5 shows two of those spectra, which correspond to SZA of about 45° and 60°. One of them presents a bias, which is not corrected here. A non-LTE model simulation has been performed for these solar illumination conditions, using the VIRA reference atmosphere, and is also presented in Figure 5. The spectral shape of the two fundamental bands of the main CO2 isotope can be discerned, by comparing with the non-LTE model simulation. The data are very noisy for further conclusions at this stage. Hopefully, a more extended data set of V-H spectra will be obtained during the rest of the mission, and a statistically sounded comparison with the theoretical prediction will be a good test for non-LTE models. This 2.7 mm non-LTE emission has also been detected in the V-M signal, as we discuss in section 3.2. 2.3.3. CO 4.7 mm Emission [25] Figures 6 and 7 illustrate a similar study for the 4.7-mm spectral region which, according to LVEA is dominated by non-LTE emissions of CO in the upper mesosphere and lower thermosphere. The emission observed is lower than the CO2 emission, and the number of V-H CO spectra with good signal to noise ratio is much smaller. About 20 spectra, out of the 145 available at the limb during daytime, show a clear signal. Many of them show a small signal above the noise level or above the daylight scattered by the aerosols and clouds. This is why it is difficult to detect this non-LTE
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Figure 5. Two selected spectra at 2.7 mm observed by VIRTIS-H (thin lines) compared to a non-LTE model simulation (thick line). The upper spectrum is from orbit 88, qube 2 at an altitude of about 107 km and SZA about 60°. The lower spectrum is from orbit 96, qube 20 at an altitude of 118 km and SZA about 45°. The theoretical simulation has been performed at 110 km and for similar solar illumination conditions. No shift has been applied to the observed data. See section 2.3.2. emission below about 70 km. The highest altitude where the emission can be clearly identified is about 120 km. [26] Three of the most interesting spectra are shown in Figure 6 (top). The two spectra with the strongest emission correspond to two observations at approximately the same tangent height, around 85 km, but for SZA of 60° and 76°. The third one, shifted downward for clarity, is the spectrum at the highest altitude available in the data set explored so far, about 120 km. In spite of the noise level, very noticeable in the high-altitude spectrum, and which amounts to 5000 mW m 2 sr 1 mm 1, in all of them we can see two branches of what seems a typical vibrational-rotational band. The simulations of LVEA for the CO(1-0), or fundamental band (FB), of the main isotope showed such structure as well. However, their prediction has a minimum emission at the FB center, at 4.67 mm, while the spectra in Figure 6 present their central minimum around 4.73 mm. This points to a different band of CO, the CO(2-1), or first hot (FH), of the main isotope, as responsible for these data. Such emission was not simulated by LVEA, who focused their analysis in the FB. These V-H measurements indicate that, if the fundamental transition is present, its intensity is smaller than the first hot. A similar effect was observed with ground-based measurements at much higher spectra resolution [see Crovisier et al., 2006]. [27] The spectral resolution of V-H allows to identify individual lines of the P and R branches of the FH CO band, and this is shown Figure 6 (bottom). A number of vertical
lines are drawn at the wavelengths of the spectral lines centers of the two bands, according to HITRAN [Rothman et al., 2005]. For clarity, we focus on the spectral interval occupied by the P branch of the FH band. We can identify most of the lines of the P branch, up to P20 at least. It is interesting to notice, however, that in the center of the FH P branch, at 4.727 mm, a weak emission line is observed, which coincides with the line P7 of the CO fundamental band. No FH lines are present there. Moreover, all the 5 lines observed around this wavelength are better fit by the P5 –P9 lines of the FB than by the FH band. In other words, the emission from the FB seems to be also detected in the central portion of the FH band. The absence of FH lines may be due to the smaller strength of the lines with low rotational number. This is a tentative explanation which requires careful quantitative modeling. [28] As previously done by Crovisier et al. [2006], it might be possible to derive temperatures from this emission after a correct model fit. This is not possible with the model of LVEA in this moment, but we have attempted to estimate the rotational temperature from some VIRTIS data. First, we assumed that the rotational levels of CO are in thermal equilibrium at the pressures of relevance. Then, the usual expression for the rotational temperature, given by the ratio of line intensities (see an example for the Venus atmosphere by Deming et al. [1983]), was applied to the spectrum shown in Figure 6 (bottom). A selection of lines from the P branch of the FH band is required. The criteria we
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Figure 6. (top) Selection of three V-H spectra at 4.7 mm. The two closer ones correspond to about 85 km and to SZA of 60° (solid line) and 76° (dashed line). The lowest spectrum, shifted 20,000 radiance units downward for clarity, corresponds to a tangent altitude of 118 km and SZA of 46°. (bottom) Zoom of the SZA 60° spectrum, with lines marking the positions of the Hitran line centers for the fundamental (dashed marks) and first hot (solid marks) bands of CO. followed were a small overlapping with the FB lines and a peak emission well above the noise levels. This method rules out high and low J lines. Also, adjacent rotational lines should be avoided in order to reduce numerical error. For this spectrum, the best subset of lines that we found are those between P3 and P8. By forming pairs between the more extreme ones, P3, P4 and P7, P8, several values of rotational temperature were found, all around 237 K, but with uncertainties of 50 K or larger. This large error is due to the large sensitivity to the peak intensity in the center of the lines. This intensity value is uncertain because it is not directly available from the data, given the V-H spectral resolution and sampling, but obtained by an extrapolation to the radiance values in the centers of the selected lines. Notice that the P branch of this spectrum apparently has a maximum at P9 and a local minimum at P6; we have excluded these two lines in this analysis since they seem specially affected by the V-H sampling and noise. This error in the lines intensities can be reduced in the future, by using a non-LTE model simulation for the FH band, as was done by Crovisier et al. [2006].
[29] Regarding the vertical variation, profiles can be constructed from the two dozens of spectra with good S/N ratio. They are shown in Figure 7 at 4.678 mm. The data have a clear dispersion, and in contrast to the model prediction for the FB, it does not show any clear peak: the plateau between 95 and 120 km, at about 6000 mW m 2 sr 1 mm 1 is very close to the noise level. As it happened with the CO2 data in Figure 4, the small number of points does not allow to discern a clear SZA variation. 2.4. VIRTIS-M Maps During Periapsis [30] In this section we present measurements of V-M from the periapsis, where the maximum spatial resolution can be achieved, as shown in Figure 2. In the Spectral Mode, near periapsis, the V-M slit is fixed, no scanning is performed, and a row of pixels is taken. Therefore, the mapping is composed of a set of successive ‘‘V-M slit frames,’’ each one taken about 10 s apart, while the spacecraft is moving along the orbit. [31] This is illustrated in Figure 8, where we plotted a map of CO2 radiances at 4.32 mm, in the center of the strong
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Figure 7. Vertical distribution or V-H radiances at a central wavelength of the 4.7-mm band of CO. Orbits and symbols as in Figure 3. See section 2.3.3. CO2 bands system. The vertical lines correspond to the actual pointing of each pixel, similarly to Figure 2. We will build similar maps with V-M data but from apoapsis observations, in the next section. [32] Two strong variations are observed in the radiances, one along the vertical axis and the other along the horizontal, and both can be explained by the non-LTE model.
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Considering first the vertical, a peak emission is observed, in the lower thermosphere, 110– 120 km. As the SZA varies along the track (from 70° to nighttime), the peak intensity, and the whole profile, decrease. This is predicted by the non-LTE model and this data set offers a perfect benchmark for a quantitative validation of the model; a work which is in progress. [33] This is the first time that maps of IR radiation with this geometrical resolution are obtained in the Venus upper atmosphere. In order to illustrate more clearly such a resolution, we constructed a vertical profile with data from a single slit view, that indicated with an arrow in Figure 8. They are shown in Figure 9 (right), at three wavelengths. A log scale is used for the radiances, to confirm that a good signal to noise is achieved even at 160 km from one single spectrum. These data correspond to SZA 75°– 80°, a value typical of daytime observations at high latitudes. The altitude of the peak is maximum in the center of the band, at 4.32 mm, and it is lower in the wings, at 4.25 and 4.40 mm, according to model expectations. Figure 9 (left) shows three V-M spectra at three tangent altitudes. The spectral resolution is lower than V-H, but the change in the overall shape of the band with altitude is clearly observed, as discussed in the previous section. We have also tried to build up maps of radiances in the 2.7-mm and 4.7-mm region. Unfortunately such V-M maps are very noisy at all altitudes, contain a very strong component of scattering up to 90 km, and are not further studied here. 2.5. VIRTIS-M Maps and Profiles During Apoapsis [34] V-M observations from the apocenter of the Venus Express orbit are good examples to show the large potential of an imaging instrument like V-M for mapping large portions of the Venus disk. However, some images also contain information about atmospheric limb emissions, and
Figure 8. Map of V-M radiances from periapsis (orbit 43) at 4.32 mm versus SZA and tangent altitude. The color code ranges from 0.03 to 0.34 mW m 2 sr 1 mm 1. Each line corresponds to a column of pixels along the V-M slit. V-H data for the same orbit are overplotted (triangles). The arrow indicates the column of pixels from which the vertical profiles in Figure 9 have been extracted. 9 of 19
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Figure 9. V-M spectra and profiles at 4.3 mm from periapsis, orbit 43. (left) Spectra at three different altitudes, as indicated. (right) Profiles at three wavelengths extracted from the pixels indicated in Figure 8. See section 2.4 for details. allowed us to extract vertical profiles with an excellent vertical resolution. 2.5.1. 2-D Images From Apoapsis [35] Figure 10 shows a map of V-M radiances at 4.32 mm, built up with the data from three consecutive qubes or V-M images (2, 1, and 0) of orbit 25, all of them acquired from
apoapsis, at about 66,000 km from the center of Venus [Svedhem et al., 2007]. The instrument takes about 10 min to take every qube, or image, before the satellite pointing switches to the next. The three images coincide with the data qubes shown in Figure 1a. The current map looks very different for two main reasons. First, now we are plotting
Figure 10. Maps of V-M limb radiances at 4.32 mm from apoapsis, from the three qubes of orbit 25 shown in Figure 1, as a function of tangent altitude and longitude. Vertical lines indicate SZA in degrees. Color scale is given in Figure 11. 10 of 19
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Figure 11. Smiling effect in V-M images from apoapsis and its correction for two cases: (top) qube 1 of orbit 25 and (bottom) qube 0 of orbit 29. (left) Original map of radiances at 4.32 mm. (middle) Location of pixels with radiance between the ranges indicated, and fit of a parabolic line to the cloud of pixels. (right) Corrected maps. See section 2.5.2 for details. just the portion of the image which corresponds to actual limb sounding, i.e., tangent altitudes above about 60 km. And second, instead of pixel number, in Figure 10 we used altitude above the surface for the y axis, and a geographical parameter, longitude, in the x axis. Since the latitude is almost constant for all these limb data, Figure 10 represents a geometrically realistic view of such altitude-longitude cross section of the Venus upper atmosphere. [36] Although the data acquisition and the image construction are different, this map is similar to the map from periapsis shown in Figure 8. Again, we observe a strong emission in the lower thermosphere, with a peak around 110– 120 km. Also a large variation is observed in the horizontal axis. The SZA variation is, however, very small in each of the three qubes/images shown in Figure 10. Therefore, the horizontal variation may be of a different nature than that seen in the periapsis case, and might indicate a real atmospheric variability. [37] However, extreme care is required before such an interpretation. The observations are taken from very far away, in a geometry which the instrument was not optimized for limb sounding; the data might contain some geometrical effects. First of all, these maps contain what appears to be a systematic bending of the radiances in the central part of the image. This bump has also been observed at other wavelengths, qubes, and dates, and it seems to be systematic. It is specially apparent in the uppermost layers.
A second problem is shown by the sharp change in radiance between adjacent qubes. The three qubes shown here were taken with some minutes apart, and there is no obvious physical reason to expect changes of that magnitude between them. Investigation of these instrumental effects is ongoing. A likely explanation for the ‘‘smiling effect’’ is discussed in the next section. A probable cause for the second problem could be some flat field effect not yet identified. 2.5.2. Geometrical Correction of V-M Images [38] The apparent bending observed in Figure 10 is of the order of 20 km, precisely the vertical resolution per pixel from the apoapsis, in other words, this error is within the size of one pixel. Although we cannot rule out small variations of the spacecraft pointing, the effect may be due to a slight bending of the V-M slit. This ‘‘smiling effect,’’ which cannot be avoided, is of special importance for this limb sounding. Fortunately, some correction is possible, precisely by using these limb non-LTE emissions. We basically used two theoretical predictions of these emissions: first their strong decrease with tangent altitude, and second their variation with SZA, as explained below. The idea is to fit the bending observed with a simple function and then use that function to replot the radiances. We illustrate two examples of such a correction in Figure 11. [39] Figures 11a and 11d show two different V-M images, those for orbit 25 (15 May 2006), qube 1 and orbit 29
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Figure 12. Vertical profiles of V-M radiance at (right) 4.33 mm after combining the (left) rows of pixels. (top) Qube 2 of orbit 25. (bottom) Qube 1. (19 May 2006), qube 2; these are the original maps to transform. The correction has to be applied to them on a pixel basis, as we will assume it to be symmetric around the central pixel of the frame 128. Then we have selected all those pixels with radiances within a small bracket of values; they are marked with crosses in Figures 11b and 11e. According to the non-LTE model, if the atmospheric densities and the SZA illumination do not change very much, a given limb radiance corresponds to a given tangent altitude. Actually, this is the reason why we discovered this smiling effect. We can observe how the crosses marking the 0.15 – 0.20 radiance interval indicate a bending approximately symmetric around pixel 128. We fit a simple parabolic function to this cloud of pixels, and used it to correct for the tangent altitude of all the 256 pixels in each row of the frame. A perfect fit cannot be obtained owing to the different sources of error, which produce some dispersion in the measurements. In addition, the cloud of data does not present a random distribution but a peculiar dispersion produced by the particular geometry of the observations, which is specific for each qube. Some atmospheric variability cannot be ruled out either. For these reasons, we aim here at a first-order correction. Let us notice, also, that it will be difficult to find a single fit function with matches all the radiance intervals of all the orbits and qubes. In our example, the radiance maps resulting after the correction are shown in Figures 11c and 11f. The bending of about 20 km is greatly reduced. This correction function introduces some uncertainty in the absolute pointing altitudes. In the center of the frame, around pixel 128, the correction is very
small, but the pointing may be inaccurate by around 5 km for the pixels at the extremes of the V-M frame. Still, we believe that the corrected images can be used for scientific studies, like atmospheric variability, possible dependence of the emissions on non-LTE parameters, etc, which would be impossible without this geometrical correction. In the VIRTIS team we are considering the implementation of this procedure in future operational processing. 2.5.3. Vertical Profiles From V-M 2-D Images [40] Lets analyze now the vertical profiles of radiance from the limb that can be obtained from these V-M images at the apoapsis. As an example, Figure 12 shows two vertical profiles extracted from two different data qubes, already shown in Figure 1, qubes 2 and 1 of orbit 25. The profiles, at the wavelength of 4.32 mm, are shown on the right, while on the left we plot a small portion of the qube/ image and marked the precise pixels used to build the profiles. The portion of the Venus disk selected corresponds to an almost constant solar illumination, with a small change of SZA, from 10° to 20°. To build up the profiles, we considered two adjacent sets of 10 pixels, taken at the same time, from two adjacent rows of the V-M frame. The selected pixels are pointing at varying altitudes, from the planet’s surface up to the Venus thermosphere, with a vertical resolution around 15 km, and symbols in Figure 12 just point to the center of each pixel. If the orientation of the limb of Venus is optimum, as we can see in Figures 12a and 12b, a very good altitude resolution can be achieved by mixing adjacent pixels; otherwise, the gain in vertical resolution is lower or nonexistent. The improvement of
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Figure 13. (top) Vertical profiles of V-M radiances at four wavelengths for qube 2, orbit 25. (bottom) Spectra of the 20 pixels composing the vertical profiles shown on the top; they belong to 2 adjacent rows of pixels, marked with stars and circles as in Figure 12. Four triangles mark the wavelengths of the four vertical profiles. the vertical resolution is at the expense of losing horizontal resolution, obviously. [41] Other four vertical profiles are shown in Figure 13 (top), obtained from the same rows of pixels in the top of Figure 12 but at four different wavelengths in the 4.3-mm region. The wavelength values are marked in Figure 13 (bottom), which shows the 20 spectra considered. We can get a lot of useful information about the non-LTE emissions from an image like this, since it combines spectral and vertical variations. For instance, the peak altitude is shown to vary with wavelength, being higher at 4.29 mm or 4.33 mm (see Figure 12) than in the wings of the band. This was predicted by the non-LTE model, and was observed also in periapsis observations. Although the qualitative behavior seems to agree with the model, the vertical profile at 4.45 mm does not coincide with the model, but this suggests a much lower emission compared with the center of the band. We think this may be solved by increasing the number of weak CO2 bands in the non-LTE model, a task for a forthcoming revision of the non-LTE model that we are currently undertaking. [42] A systematic problem in V-M data can be noticed in Figure 13 (bottom). There seems to be an oscillation in the spectral domain, with high and low values of intensity from one spectral point to the next. We checked that this ‘‘oddeven effect’’ is present at all VM near-IR wavelengths, from 2 to 5 mm, and it amounts to about 0.05 radiance units in Figure 13, or 50,000 mW m 2 sr 1 mm 1, which is about 10 times larger than the measurement noise. The technical cause of the odd-even seems to be an asymmetry of the
Read-Out Integrated Circuit (ROIC) of the Focal Plane Array (FPA) of the instrument; see Coradini et al. [1998] for a more detailed description of the instrument. In particular, there is an asymmetry in the clock feedthrough and the unit cells between pixels in odd columns and pixels in even columns, which causes some differences in gain and offset. Most part of this effect is corrected through the radiometric calibration where the responsivity of each pixel is taken into account. However, there seems to be also a problem of the offset/reset being dependent with the signal, which makes the difference of the odd-even effect bigger for extreme conditions, which is the case of very high signals or very low exposure times. During periapsis, for example, the exposure time is about 10 times larger than in apoapsis. This is the reason why the odd-even effect is much smaller (at or below the noise level) in Figure 9. Let us recall that odd-even effects are common in imaging spectrometers, in the spectral or spatial domains [Moutou et al., 2003; Siebenmorgen et al., 2007]. We will return to this point in the next section.
3. Validation Analysis [43] Validation analysis traditionally comprises direct comparisons with independent observations. This is not possible with our VIRTIS limb data, although eventually, some indirect comparisons with related measurements by other Venus Express instrument, like neutral densities from VERA and SPICAV, will be possible. At this stage, however, our analysis focused on other aspects. These include (1) internal consistency tests and comparison with averages,
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Figure 14. V-H and V-M noise and bias in a selection of spectra around 4.3 mm from periapsis, at a tangent altitude of 118 km. Radiance units are mW m 2 sr 1 mm 1. See section 3.1 for further details. (2) estimation of noise/data quality, (3) cross correlation between the two VIRTIS signals, (4) qualitative comparisons with theoretical non-LTE models, and (5) examination of the repeatability and variability of the measurements. The goal is to characterize the behavior of the VIRTIS instrument, to confirm that the measurements are physically meaningful, and to detect potential bias or systematic errors. The most important of these analysis, and their conclusions, are presented next. 3.1. Noise and Bias in V-H and V-M [44] One of our basic objectives was to evaluate the measurement noise in orbit, in the spectral ranges and emissions of our interest, to compare it with the nominal one obtained by ground calibration, and at the same time, to analyze possible systematic effects. A first step was the examination of individual spectra for the orbits with limb data. A second task was to average homogeneous sets of data, considering appropriate boxes in SZA and tangent altitude, and to calculate their mean radiance and dispersion. [45] For the evaluation of the noise level and the bias in the 4.3-mm region, we looked at the two wings of this CO2 system of bands, shortward and longward of 4.20 and 4.50 mm, and at altitudes where no significant emission is expected there. The mean radiance of each spectrum can be considered as the bias, and its standard deviation (SD) as an estimation of the noise. An example for orbit 43 (periapsis) is shown in Figure 14, where two individual spectra are shown, one from V-H and a nearby one from V-M. The tangent altitude is around 120 km. [46] Regarding V-H, the noise values obtained for this particular example are similar in the left and right wings of the band, and slightly lower than the nominal noise of 5000 mW m 2 sr 1 mm 1. We have evaluated these values
for a diverse set of spectra, for different orbits, solar illumination conditions and tangent altitudes, and have found that they are basically constant up to orbit 47 (7 June 2006) and decrease to about half those values after orbit 79 (8 July 2006). This is probably related to the spacecraft thermal evolution along the mission, with higher noise levels at the beginning of the mission, when the cold box temperature was higher. [47] Regarding the V-M data, the noise in this example is also lower than the nominal value, although there is a significant difference between the left and right wings of the band. Regarding its bias, the availability of a large number of spectra from V-M at a given position, allows us to estimate it for a close subset of pixels. In Figure 14 we added another spectrum, that obtained by averaging 25 V-M pixels/spectra within a small range of tangent altitudes (118 –120 km) and SZA (78° –81°) around the V-M spectrum selected. The mean value coincides well with the bias obtained from one single spectrum. [48] Similar noise and bias analyses for V-M at apoapsis are difficult at present owing to the large odd-even effect mentioned above, much larger than those parameters. In fact, the averaged V-M spectrum in Figure 14 shows more clearly such oscillation in the intensity than one single spectrum. This effect has a magnitude (amplitude between peaks) of about 2000 mW m 2 sr 1 mm 1, of the order of the noise. The V-M calibration group of the VIRTIS team is currently working on this point, in order to reduce its impact on apoapsis data. 3.2. Correlation Between V-H and V-M Spectra [49] There is a small number of observations in the Spectral Mode which correspond to the special Tangential Limb sounding, with simultaneous measurements of V-H
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Figure 15. Correlation between V-M and V-H spectra taken during periapsis, orbit 43. (right) Location in tangent altitude and SZA, with approximate size of the V-H and V-M field of view. (left) Spectra of the 15 selected V-M pixels, those with SZA of 78.2°, 79.8°, and 81.3°, and of the V-H spectrum (convolved to V-M spectral resolution). Differences between V-H and V-M spectra are also plotted below. See section 3.2 for further details. and V-M. The proximity to the planet during periapsis, allows for a precise geographical collocation of these profiles, and therefore, permits a specially appropriate comparison between them. Orbit 43 offers one example, and Figure 8 shows how close V-M and V-H measurements can be. Figure 15 shows a number of VM spectra close to one particular V-H spectrum, obtained at that orbit, around SZA = 80° and 119 km tangent altitude. Five adjacent slit images were considered, with five pixels on each. The vertical size of the pixel’s FOV, about 300 m for V-M and 2 km for V-H, is shown on the right. On the left we plot 15 of the VM spectra, those from the groups at SZA of 78.2°, 79.9°, and 81.3°, and compare them with the V-H spectrum. This was degraded/convolved to the lower spectral resolution of V-M for the comparison. The shape of the whole CO2 band is very well reproduced by all of them. We think the overall comparison seems fairly good, given the dispersion among the 15 V-M spectra plotted. These present, first of all, a small dispersion within each group, which is similar or slightly larger than typical noise and bias levels. In addition there is a small variation between the three groups, significantly larger than the noise, which is surely due to the small SZA difference between them; the radiance at larger SZA being smaller, as expected. Comparing the V-H spectrum with its closest V-M group of spectra, there is a small shift between them, of about 40,000 mW m 2 sr 1 mm 1 around the peak at 4.33 mm, larger than the noise and bias of both, V-H and V-M. However, this difference is only a factor 2 larger than the SD of this particular group of V-M data, and equivalent to a small fraction of a degree in SZA. This gives us a good level of confidence on both signals.
[50] The emission at 2.7 mm can also be used to test the V-H and V-M agreement. Although the number of spectra with good signal is very limited, we found some spectra in the orbit 43 which are useful for the present discussion. They are shown in Figure 16 and correspond to SZA = 80° approximately. Figure 16 (top) shows vertical profiles at three wavelengths in the 2.7 mm spectral region and show a neat altitude variation, with a clear peak around 115– 120 km. This was predicted by the LVEA model. Figure 16 (bottom) shows the mean and standard deviation of a set of 15 V-M spectra, at altitudes around 120 km. The error bars of individual V-H and V-M are slightly larger than the nominal values, and they have been omitted in Figure 16. The agreement between V-H (convolved to V-M resolution) and V-M is good within error bars; there may be a slight mismatch around 2.8 mm but the data seem noisier there. Moreover, this emission is lower and noisier than that shown in Figure 5, at all wavelengths, since the solar excitation is lower now. We performed a simulation with the non-LTE model, also shown in Figure 16, for the same solar illumination conditions. The agreement is very encouraging, although the model seems to overestimate the radiance; this is due to the specific atmosphere structure assumed in the simulation, and therefore is of no relevance in this analysis. [51] Related to this correlation, we investigated the correlation between different V-H spectra, wherever they can be collocated. The amount of V-H spectra is not large, but we identified a couple of interesting correspondences between spectra with good S/N ratio. On Figure 17 (left), they correspond to four consecutive V-H measurements (acquisition 84 to 87) obtained at the closest approach to Venus
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Figure 16. (top) Vertical profiles of V-H and V-M radiances at three wavelengths in the 2.7 mm for qube 0, orbit 43 and SZA around 80°. Triangles and crosses mark V-H and V-M data, respectively. (bottom) One V-H spectrum, convolved to V-M resolution (solid line) and comparison with the mean of 15 V-M spectra (dashed lines), from the data composing the vertical profile in the top, at a tangent altitude of 120 km. The dashed-dotted line represents a model simulation. Radiance units are mW m 2 sr 1 mm 1 in both the top and bottom. Three triangles in the bottom indicate the wavelengths of the three vertical profiles. during orbit 43, and point to the lowest tangent altitudes for this orbit, around 99 km (as indicated in the picture) and around SZA = 83° (see Figure 8). The geometrical distance along the track of the satellite between the spectra 84 and 87 is about 250 km, and they correspond to latitudes of 73.0° and 71.8°, longitudes of 302° and 307°, respectively, and a difference in SZA of 1.4°. The four spectra are very similar, with differences of the order of the noise at almost all wavelengths, except in the central region. Around 4.3 mm there is a clear trend, from 84 to 87, the last one being about 30,000 mW m 2 sr 1 mm 1 lower than the former. We think, again, that this is due to SZA changes; close to the terminator small changes in SZA produce noticeable variations in the solar pumping of the CO2 vibrational states. This reflects small changes in the atmospheric structure between those four locations. [52] Figure 17 (right) presents another two V-H spectra, very closely located to each other, both at a latitude of 77.27°S and longitudes of 257.8°E and 260.8°E, respectively, but obtained 2 days apart, on orbits 40 and 42. Their tangent altitudes and SZA are also very close, at 108.4 and 107.8 km, and 82.7° and 82.3°, respectively. Their noise and bias levels are below 5000 mW m 2 sr 1 mm 1 and 8000 mW m 2 sr 1 mm 1, respectively. It is surprising the excellent agreement between them in the overall shape and in the wings of the band; there is only a small difference, in
the central portions of the CO2 emission, with the spectrum of orbit 42 slightly larger than the other and the noise. Again, the small SZA differences is a possible candidate for the observed difference. This points, once more, to a relatively calm and stable atmospheric structure, fully repeatable in times scales of 2 days, at least in this case. It would be interesting to explore further coincidences of this type in future measurements of V-H, since the number of them available to date is small (see Figure 3). 3.3. Comparisons With a Non-LTE Model [53] As discussed above, we have used in this work the non-LTE model of the Venus atmosphere developed at the Instituto de Astrof ´ısica de Andalucı´a/CSIC, in Granada [Rolda´n et al., 2000; Lo´pez-Valverde et al., 2007], to perform a small number of simulations. These simulations and those reported by LVEA have guided most of the analysis presented in previous sections. From these comparisons we found a good overall agreement with the predictions of the spectral shape of V-H and V-M data (Figures 4, 9, and 17), and with the altitude profiles of the radiances in the central part of the 4.3-mm band (Figures 4 and 9). Also the SZA variation of the peak radiance (Figures 8 and 15) qualitatively agrees with expectations. The vertical profiles of the CO2 emission shown in Figures 9 and 13 do agree qualitatively with those of LVEA but a quantitative fit
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Figure 17. Geographical coincidences between different V-H spectra. (left) Four spectra from orbit 43, around 99 km. (right) Example of two spectra taken at very close points but 2 days apart. Variations attributable to SZA changes. See section 3.2 for details. would be desirable and future work should focus on this aspect and exploit such fit, once the weak bands required are implemented into the model. Other non-LTE emissions predicted by LVEA have also been examined, detected and compared with the model. Figures 5 and 16 show simulations of the CO2 2.7 mm bands for a given reference atmosphere; they indicate a good agreement in the spectral shape of the bands, and for both V-H and V-M spectral resolutions. Although these data are noisy, the main features are captured by the model, and the vertical profiles are also in agreement, with similar altitudes of the peak emission in the data and in the model. The absolute emission, as it happens with the 4.3 mm bands, is not considered in this comparison, since it depends on the actual atmospheric structure assumed in the non-LTE model. The third nonLTE emission detected and studied in this work, that of CO, contains a first hot component not simulated yet by the non-LTE model. [54] Let us highlight here some other results which, either were not predicted or seem to pose interesting challenges to the theoretical modeling. One of them is the intensity in the longer wavelength wing of the 4.3 mm band, where contributions from weak, isotopic and hot transitions are expected. The predictions by LVEA produce a relatively minor contribution of that band’s wings compared to the central band, but the VIRTIS spectra shown in Figures 4 and 9 indicate a more significant contribution. The model surely needs a larger number of CO2 bands in this spectral region. Such a model extension is underway. [55] Certainly, another extension of the non-LTE model is needed to include the first hot band of carbon monoxide, CO(2-1), in order to study the spectra shown in Figure 6.
This is also a minor band compared to the strongest fundamental transition, but it seems to be highly excited by direct solar radiation, like the weaker transitions of CO2 discussed above. As discussed in section 2.3.3, a proper simulation of the CO emission lines will be very valuable for the derivation of rotational temperatures. These could be obtained from this band’s structure, but only after fitting a theoretical model which permits small radiance errors in the line peaks. This is the only molecular band detected by VIRTIS in this infrared portion of the spectrum with well isolated lines, and is therefore a very interesting data set.
4. Conclusions and Future Work [56] VIRTIS on Venus Express is an innovative instrument which, combining imaging and spectroscopy in the near infrared, is obtaining very exciting new data on the upper atmosphere of Venus. We have analyzed the limb measurements taken by VIRTIS in the regions from 4 mm to 5 mm and around 2.7mm from both, the periapsis and the apoapsis of the Venus Express orbit, and in the two signals, V-H at high spectral resolution, and V-M at a lower resolution. We have focused in this work on a subset of VIRTIS data, which includes limb measurements with V-H taken at the periapsis from orbits 23 (13 May 2006) to 127 (25 August 2006), and a small sample of V-M data qubes between orbits 23 (13 May 2006) and 295 (10 February 2007). Extension of the analysis presented here to the rest of the data is ongoing. [57] We clearly identified the strong emission by CO2 at 4.3 mm in the upper mesosphere and lower thermosphere, and of the first hot band of CO at 4.7 mm in the upper
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mesosphere. Individual spectra show good signal to noise up to above 160 km in the case of CO2, and up to about 120 km in the case of CO. Single spectra show noise values slightly below the value expected from ground calibration, and a small bias too. We also showed that the V-H spectral resolution is good enough for a number of tests of our nonLTE models. The spectral resolution of V-M is lower, however the vertical resolution is much better from V-M data. During periapsis this is typically 300 m for a single row of pixels, while from apoapsis, the grouping of V-M adjacent rows of data permits a resolution at least two to three times better than the FOV of individual pixels, depending on the geometry of the V-M images and on the relative orientation of the limb of Venus. These observations characterize the Venus atmosphere directly, since they will permit to calculate densities and temperatures from measurements, and also indirectly, because they can give information about the collisional processes, energy transfer and the upper atmosphere energy balance, after a more detailed comparison with a non-LTE model. [58] The variability observed by VIRTIS and mentioned previously by Drossart et al. [2007a] have been confirmed and analyzed here for the first time with this new and extended data set. The maps studied in this paper permit to describe altitude variations at several wavelengths, variations with SZA, and the detailed spectral shape of the CO and CO2 bands. Actually, the good agreement between model and data, on the basic non-LTE features, indicates that our current understanding of the physical processes in the upper atmosphere is essentially correct. All these results will be perfect tests for future non-LTE models, and suggestions for more detailed future comparisons are given in section 3.3. We foresee that such detailed comparisons, specially of the shape of the V-H emissions, can give us information about non-LTE collisional parameters in the upper mesosphere and lower thermosphere, where these processes are specially relevant. Also, comparisons between spectra at similar conditions but obtained with a span of 2 days, and between data more than 200 km apart, show very good matches, which indicates a good behavior of VIRTIS. When applied to a future and extended data set, similar comparisons may supply interesting information about atmospheric variability in the lower thermosphere. The high density of points in the limb in some maps during periapsis offer excellent chances for derivation of density maps in the thermosphere. Such density profiles are foreseen to be obtained by a rigorous ‘‘retrieval’’ process which uses inversion techniques under non-LTE conditions, an exercise similar to other upper atmosphere experiment on Earth, like the MIPAS/Envisat data analysis [Funke et al., 2005]. Densities of CO would be derived in a limited altitude range (90 – 120 km) owing to the low signal of the CO 4.7 mm emission (see section 2.3.3). However, densities of CO2 can be derived in principle in a wider altitude range. Nevertheless, the accuracy of such derivation will depend on a sensitivity study to be performed with the revised non-LTE model. Limb views from V-M during apoapsis should also be very useful for density retrievals, although they are affected at present from a ‘‘smiling effect’’ which deforms the actual emission field by about 20 km, approximately the size of the projected FOV of one pixel. We showed how the non-LTE model has been essential to detect this effect which could be reduced in
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future processing thanks to a correction technique using these CO2 non-LTE emission. These V-M data from apoapsis are also affected by a large odd-even effect, about 10 times larger than the noise level. During periapsis this effect is below noise due to the much larger exposure time there. This aspect requires further work within the VIRTIS team in the future. [59] The study of the first hot CO(2-1) emission is also very promising, since the rotational structure is resolved and it can therefore be used to derive rotational temperatures in the upper mesosphere. This will be confined to daytime and to the altitude range 90– 120 km, where such signal is strongest. This derivation does not require a full non-LTE retrieval but a fit of the shape of the CO emission band with the model, as explained in section 2.3.3. At present, uncertainties in individual lines of a CO spectrum produce very large temperature errors. Comparisons with temperatures obtained from ground-based measurements of these bands [Crovisier et al., 2006] will also be interesting, as a validation exercise of these data and in order to future check the non-LTE models. [60] VIRTIS/Venus Express will continue acquiring new data at least until 2009. The exploration of the new data and their analysis will continue, hopefully with a version of the non-LTE model extended to weaker bands, and with better understanding of some of the current uncertainties in the data that we highlighted here. The application of retrieval techniques and systematic search for atmospheric variability, makes the VIRTIS observations a very promising data set for studying the non-LTE processes and the structure of the upper atmosphere of Venus. [61] Acknowledgments. We thank the space agencies ASI and CNES for their support. The work of the IAA-CSIC team has been carried out under project ESP2004-01556 and was partially funded by the project AYA2008-03498/ESP of the Spanish Ministery of Science and Innovation and by EC-FEDER funds. GG has been partially supported by a fellowship from Spanish National Research Council (CSIC-JAE Program).
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Moutou, C., A. Coustenis, J. Schneider, D. Queloz, and M. Mayor (2003), Searching for helium in the exosphere of HD 209458b, Astron. Astrophys., 405, 341 – 348. Piccioni, G., et al. (2006), VIRTIS: The Visible and Infrared Imaging Spectrometer, Eur. Space Agency Spec. Publ., ESA-SP 1295, 1 – 27. Rolda´n, C., M. A. Lo´pez-Valverde, M. Lo´pez-Puertas, and D. P. Edwards (2000), Non-LTE infrared emissions of CO2 in the atmosphere of Venus, Icarus, 147, 11 – 25, doi:10.1006/icar.2000.6432. Roos-Serote, M., P. Drossart, Th. Encrenaz, E. Lellouch, R. W. Carlson, K. H. Baines, F. W. Taylor, and S. B. Calcutt (1995), The thermal structure and dynamics of the atmosphere of Venus between 70 and 90 km from the Galileo-NIMS spectra, Icarus, 114, 300 – 309, doi:10.1006/icar.1995.1063. Rothman, L., et al. (2005), The Hitran 2004 molecular spectroscopic database, J. Quant. Spectrosc. Radiat. Transfer, 96, 139 – 204. Siebenmorgen, R., et al. (2007), Very large telescope paranal science operations CRIRES user manual, Rep. VLT-MAN-ESO-14500-3486, 56 pp., Eur. South. Obs., Garching, Germany. Stepanova, G. I., and G. M. Shved (1985), Radiation transfer in the 4.3-mm CO2 band and the 4.7-mm CO band in the atmospheres of Venus and Mars with violation of LTE: Populations of vibrational states, Sov. Astron., Engl. Transl., 29, 248 – 422. Svedhem, H., et al. (2007), Venus Express: The first European mission to Venus, Planet. Space Sci., 55, 1636 – 1652, doi:10.1016/j.pss.2007.01.013. Taylor, F. W., et al. (1980), Structure and meteorology of the middle atmosphere of Venus infrared remote sensing from the Pioneer orbiter, J. Geophys. Res., 85, 7963 – 8006. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017. A. Cardesı´n Moinelo and G. Piccioni, IASF, INAF, via del Fosso del Cavaliere 100, I-00133 Rome, Italy. P. Drossart and S. Erard, Observatoire de Paris, F-92195 Meudon, France. G. Gilli and M. A. Lo´pez-Valverde, Instituto de Astrofı´sica de Andalucı´a, CSIC, Camino Bajo de Huetor 50, E-18008 Granada, Spain. (
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Near-IR oxygen nightglow observed by VIRTIS in the Venus upper atmosphere G. Piccioni,1 L. Zasova,2 A. Migliorini,1 P. Drossart,3 A. Shakun,2 A. Garcı´a Mun˜oz,4 F. P. Mills,4,5 and A. Cardesin-Moinelo1 Received 29 February 2008; revised 9 January 2009; accepted 17 February 2009; published 30 April 2009.
[1] We present observations of both the (0–0) and (0–1) bands at 1.27 and 1.58 mm of the O2(a1Dg X3S g ) nightglow made with the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) instrument aboard Venus Express. The observations were conducted in both nadir and limb viewing modes, the latter constituting the first systematic investigation into the vertical distribution of the volume emission rate of the infrared oxygen nightglow in Venus’ upper atmosphere. Limb measurements from 42 orbits covering the latitude range 7°S to 77°N are analyzed. The peak altitude of the volume emission rate occurs typically between 95 and 100 km, with a mean of 97.4 ± 2.5 km. The vertical profile is broader near the equator, with a full width at half maximum of 11 km, a factor 2 larger than at middle latitudes. A double peak is frequently observed, with the lower and upper peaks occurring near 96–98 km and 103–105 km, respectively. On average, the nightglow appears brightest in the vicinity of the antisolar point. This conclusion is consistent with past ground-based observations and nadir measurements by VIRTIS. We mapped the global mean O2 nightglow intensity from VIRTIS data collected during 880 orbits. Patchy features of the nightglow intensity observed in nadir view are correlated with the thermal brightness at 4.23–4.28 mm. The observed positive correlation is consistent with downwelling (upwelling) of oxygen atoms accompanying compressional heating (expansion cooling) or with modulation by gravity waves. Finally, from simultaneous measurements of the 1.27 and 1.58 mm bands, we have estimated the ratio of the transition probabilities A00/A01 to be 63 ± 8. Citation: Piccioni, G., L. Zasova, A. Migliorini, P. Drossart, A. Shakun, A. Garcı´a Mun˜oz, F. P. Mills, and A. Cardesin-Moinelo (2009), Near-IR oxygen nightglow observed by VIRTIS in the Venus upper atmosphere, J. Geophys. Res., 114, E00B38, doi:10.1029/ 2008JE003133.
1. Introduction [2] Intense nightglow emission in the (0 – 0) band of the O2 (a1Dg X3S g ) electronic transition near 1.27 mm was first detected on Venus in 1975 [Connes et al., 1979]. Subsequent ground-based observations of the airglow emission on the nightside (nightglow) have shown that the intensity and spatial distribution are highly variable on time scales of hours to days [Crisp et al., 1996; Ohtsuki et al., 2005, 2008; Bailey et al., 2008]. Nadir-looking observations by the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) on Venus Express (VEX) have provided more
1
INAF, Istituto di Astrofisica Spaziale e Fisica Cosmica, Rome, Italy. Space Research Institute of Russian Academy of Sciences, Moscow, Russia. 3 LESIA, Observatoire de Paris, UPMC, Universite´ Paris-Diderot, CNRS, Meudon, France. 4 Research School of Physics and Engineering, Australian National University, Canberra, ACT, Australia. 5 Fenner School of Environment and Society, Australian National University, Canberra, ACT, Australia. 2
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003133$09.00
detailed coverage of the morphology of the nightglow, and visual tracking of bright features has enabled the estimation of apparent motions [Hueso et al., 2008]. Limb observations by VIRTIS of the O2 nightglow [Drossart et al., 2007a; Ge´rard et al., 2008] provide a complementary view of the chemistry and dynamics of Venus’ upper atmosphere by determining the vertical distribution of the O2 nightglow. However, no systematic study of a large number of limb observations of the O2 nightglow has been published to date. In this work, we examine limb observations spanning 7°S to 77°N latitude and 2000 to 0400 local times. These observations were collected during the period from July 2006 until August 2008. This much larger set of data has allowed us to determine how the vertical distribution of O2 nightglow varies with latitude. We have also been able to make an observational estimate of the ratio of the transition probabilities for the (0 – 0) and (0– 1) bands near 1.27 and 1.58 mm, respectively. Our results provide new observational insight into the chemical and dynamical processes occurring in Venus’ upper atmosphere and new constraints on chemical transport models. [3] On Venus’ nightside, O2(a) is believed to be produced primarily via reaction (1) [e.g., Connes et al., 1979; Mills
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and Allen, 2007], where O*2 denotes an electronically excited state of O2: 2O þ M ! O*2 þ M
ð1aÞ
O*2 þ M ! O2 ðaÞ þ M
ð1bÞ
There are two recent estimates of the net yield of O2(a) from reaction (1). Crisp et al. [1996] inferred net yields of 0.63 ± 0.19 and 0.6– 0.75 for M = N2 and CO2, respectively, on the basis of laboratory studies, while Huestis [2002] recommends a net yield of 0.94 – 0.99 for M = N2 and CO2 on the basis of a combination of laboratory and atmospheric studies. The atomic oxygen consumed in reaction (1) is transported to the nightside via the upper atmospheric circulation from the dayside [Bougher and Borucki, 1994] where atomic oxygen is produced by photodissociation of CO2 at wavelengths <200 nm. Nightglow in the O2(a-X) bands results from O2 radiatively decaying from the electronically excited O2(a) state to the ground electronic state. The vertical profile of the a-X nightglow is determined by the balance between production of O2(a) via reaction (1) and its loss due to radiative decay of the O2(a) state and collisional quenching. There has been controversy over the lifetime of O2(a) against radiative decay, but recent laboratory measurements [Cheah et al., 2000; Lafferty et al., 1998; Miller et al., 2001; Newman et al., 1999; Spalek et al., 1999] appear to have converged to a consistent value. The mean lifetime we have computed from these studies, 74 ± 3 min is consistent with recent assessments [Gamache and Goldman, 2001; Slanger and Copeland, 2003]. Only an upper limit on the rate constant for collisional quenching of O2(a) by CO2 has been determined in laboratory studies [Sander et al., 2006], so coordinated limb observations of the O2(a-X) nightglow, total density, and temperature may provide the best estimate to date of this rate constant. A recent calculation estimated 80% of the O2(a) produced at 96-km altitude decay via emission in the a-X bands [Ge´rard et al., 2008]. More detailed studies of VIRTIS’ limb observations of O2(a-X) nightglow should provide further insight into both the physical and chemical conditions in the nightglow region and the processes involved. [4] The upper atmospheric circulation on Venus is a composite of two patterns: a subsolar (SS) to antisolar (AS) flow and a retrograde zonal flow. Long-term averages of NO and O2 nightglow emissions suggest that the longterm average circulation at 115– 150 km altitude is dominated by strong subsolar to antisolar flow with a weaker retrograde zonal flow [Bougher and Borucki, 1994]. Tracking of cloud features at 65– 70 km shows the cloud top circulation is dominated by the retrograde zonal flow [Gierasch et al., 1997; Markiewicz et al., 2007; Limaye, 2007; Sa´nchez-Lavega et al., 2008]. At intermediate altitudes, the relative strength of these two patterns varies with altitude and with time [Lellouch et al., 1997; Bougher et al., 2006]. Part of the temporal variability is likely due to breaking gravity waves transporting energy and momentum upward from the cloud tops [Alexander, 1992], but significant further work is required to understand and quantify
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their effects [Bougher et al., 2006]. VIRTIS’ limb observations of O2(a-X) nightglow may provide insight into the vertical and horizontal transport processes and rates, complementing the horizontal transport results obtained using nadir-looking images from VIRTIS [Hueso et al., 2008]. [5] The first reported detection of nightglow emission in the a-X (0 – 1) band of O2 near 1.58 mm on Venus was based on VIRTIS data [Piccioni et al., 2008]. This band has not been observed on Mars but has been observed on Earth [Vallance Jones and Harrison, 1958; Winick et al., 1985]. No prior observation on any planet, however, has detected simultaneously the a-X (0 – 0) and (0– 1) bands. Significant uncertainties exist in both laboratory and theoretical determinations of the relative transition probabilities for these bands, so VIRTIS’ observations of the two bands should advance our understanding of the molecular physics of the a and X states of O2.
2. Instrument Description and Observation Geometry [6] VIRTIS is an imaging spectrometer which covers the spectral range 0.27 to 5.1 mm [Piccioni et al., 2009; Drossart et al., 2007b]. It includes two spectrometers: VIRTIS-M, a mapping spectrometer with medium spectral resolution, and VIRTIS-H, an echelle spectrometer with higher spectral resolution than VIRTIS –M but no imaging capability. VIRTIS-M includes two channels: a visible channel spanning 0.27 to 1 mm and an infrared channel spanning 1 to 5.1 mm. The results reported here were obtained using the infrared channel of VIRTIS-M, with spectral sampling interval of about 10 nm, a spectral full width at half maximum (FWHM) of about 15 nm, and an instantaneous field of view (IFOV) of 250 mrad. As is typical for an imaging spectrometer, a single instantaneous acquisition by VIRTIS-M collects a frame, which is composed of 432 bands (wavelengths) for each of 256 samples (spatial pixels along the instrument slit). If the slit is oriented perpendicular to the limb of Venus, then a single limb observation will obtain a full spectrum at each of 256 tangent altitudes, so the limb profiles of the a-X (0 – 0) and (0– 1) nightglow emissions can be derived from a single frame. A scanning mirror (or the spacecraft’s motion) is used to image multiple sites with different latitudes, longitudes, and/or local times and, thus, obtain a vertical cross section of the line of sight integrated nightglow emission. A similar process is used in nadir imaging mode to obtain 256 pixel 256 pixel 432 band spectral imaging cubes. A full spatial scan across the 64-mrad field of view (FOV) of the mirror requires around 50 min with an 8-s exposure time for each frame. The mirror moves nominally by 250 mrad during the acquisition of an 8-s frame, so that the typical pixel is square, but both the mirror’s scan rate and the integration time can be varied. The orientation of the VIRTIS-M slit is fixed relative to the spacecraft, so it is not always possible to orient the VIRTIS-M slit perpendicular to Venus’ limb when conducting limb observations. The least favorable observation geometry is when the VIRTIS-M slit is tangential to Venus’ limb because then every spatial pixel samples approximately the same tangent altitude. However, even a 10° angle between the VIRTIS-M slit and Venus’ limb is sufficient to obtain a nearly complete limb profile.
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The absolute pointing accuracy for each pixel is much better than the IFOV. This is achieved thanks to the three-axis stabilized capability of the Venus Express spacecraft, supported by accurate star trackers which provide an absolute pointing accuracy better than ±35 mrad (about ±1/7 of a VIRTIS-M pixel) in the timeframe of the observation. Other observation geometries are possible [Titov et al., 2006]. [7] Limb and nadir observations provide complementary information on nightglow emissions. During limb observations, the line of sight integrated volume emission is about 50 times larger than in nadir mode, so faint emissions from minor species are more easily detected. The vertical spatial resolution (2.5 km at 10,000-km slant distance), however, limits the maximum distance from which high-quality limb observations can be made. Venus Express’ polar orbit has its pericenter at about 75°N, so almost all limb observations are made in the northern hemisphere. Nadir observations provide an extended view of the geographic distribution of nightglow emission but no information on its vertical distribution. Spatial resolution decreases when the spacecraft is further from Venus (15 km at apocenter), but areal coverage increases, so nadir observations are typically made over the southern hemisphere and areal coverage is extremely poor north of 20°N. Examples of observations in limb and nadir modes are shown in Figure 1 for orbits 76_18, 271_09, and 44_01 (prefix before the underscore is orbit number, suffix is session number, which is the period of time of continuous acquisition of data resulting in a spectral imaging cube). Figures 1a and 1b show an example of an observation made with the VIRTIS-M slit nearly perpendicular (in the central part of the scan, executed from top to bottom) and tangent to Venus’ limb, respectively. Figure 1c (right) shows an example of nadir view.
3. Data Selection [8] For this study, we have used both limb and nadir observations. The limb observations were acquired during 72 sessions from 42 orbits between July 2006 and August 2008, Table 1. Each orbit lasts 24 h and orbit number 5, the first stabilized nominal orbit, corresponds to 25 April 2006. Usually a few sessions are acquired within a single orbit. Each spectral imaging cube typically extends over 30° latitude, and the 72 sessions selected span 7°S to 77°N. These 72 sessions include all VIRTIS-M limb observations with exposure time of 8 s. Long exposure times (8 s) provide an excellent signal-to-noise ratio (SNR), typically better than 100, at wavelengths shorter than 4 mm, which enabled derivation of well defined vertical profiles, including the high-altitude tail, for fainter nightglow emissions, such as the a-X (0– 1) band at 1.58 mm. The data from the 72 sessions selected have a vertical spatial resolution of about 2 km. The limb profiles of the O2 emission are obtained for each 5° latitude interval as the moving average within 1 km of altitude (in vertical direction). Vertical profiles are retrieved from the observed limb profiles using an onion peeling technique. Most of the limb mode observations in Table 1 had a limb profile with a single peak. These observations were used for retrieval and study of the emission layer’s characteristics. In Table 1 we report only the average of the parameters for each session. The three limb profiles in Table 1, selected as an example among the
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many others with double peaks but not reported, are discussed below but were not used for our statistical analyses. [9] The nadir observations that were examined in greatest detail in this study are reported in Table 1. For simplicity here we call ‘‘nadir observations’’ all those where the instrument line of sight (LOS) intercepts the solid body of the planet, although the term ‘‘nadir’’ is more appropriate for perpendicular view only. All have exposure times of 3 to 18 s, since the a-X(0– 0) band at 1.27 mm is sufficiently bright to provide enough signal for mapping its geographic distribution with these exposure times. The spatial resolution for the selected nadir observations is typically about 10 km. [10] The global nadir view map for nightglow was constructed using all data collected during 880 orbits between 14 May 2006 and 14 September 2008 with emergence angle less than 80° (0° = nadir). The spanned period corresponds to about 4 Venusian years. The oxygen nightglow map was constructed from observations with exposure times of at least 0.1 s.
4. Data Processing and Separation of Thermal and Nightglow Emissions [11] Nightside infrared spectra of Venus contain nightglow emission from multiple species and thermal emission, which originates in the lower atmosphere [Allen and Crawford, 1984]. In the spectral regions where absorption by CO2 is small, the thermal emission can be observed outside Venus’ atmosphere after having undergone multiple scattering within Venus’ cloud layers [Kamp et al., 1988]. One window in the CO2 absorption spectrum lies near 1.27 mm [Meadows and Crisp, 1996], so emission originating from below 20 km altitude [Tsang et al., 2008] and scattered by higher-altitude clouds overlaps spectrally and, for nadir observations, spatially with the a-X(0 – 0) nightglow emission. Although the peak of the a-X(0 – 0) nightglow emission occurs at a shorter wavelength than the maximum of the thermal emission, it is not possible to spectrally resolve and quantitatively separate the two sources with VIRTIS-M’s spectral resolution. However, nearby spectral windows, at 1.18 and 1.74 mm, provide information on the thermal emission from the lower atmosphere that may be used to assess the contribution of thermal emission to the observed 1.27 mm radiance and to empirically remove the thermal emission contribution when significant. The nightside radiance at 1.18 mm originates from the surface and the lowest 20 km of the atmosphere [Meadows and Crisp, 1996]. The 1.74 mm radiance originates from near 20– 28 km [Meadows and Crisp, 1996]. Given existing uncertainties in the CO2 spectrum near 1.27 mm [Tsang et al., 2008], an empirical approach for removing the thermal emission contribution at 1.27 mm is probably as accurate as more detailed radiative transfer calculations. [12] For limb observations, as shown in Figure 2a, the thermal emission at 1.18 and 1.74 mm is much smaller than the radiance at 1.27 mm between 90 and 103 km. The difference is attributed to the nightglow and therefore, given the expected correlation among these bands for the thermal emission, it is safe to neglect the thermal emission at 1.27 mm in this altitude range. Some typical limb profiles of the oxygen nightglow intensity at 1.27 mm are shown in
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Figure 1. Limb and nadir view modes. (a) Orbit 76_18, case of limb observation with slit (parallel to the horizontal direction) almost perpendicular to the limb in the central part of the image. The image of the oxygen nightglow emission at 1.27 mm and the images in 1.18 and 1.74 mm (thermal emission of the lower atmosphere scattered by the clouds) windows are shown. The isolines are the tangent altitude above the Venus ellipsoid. Scattered thermal emission extends up to altitudes exceeding 80 km. (b) Orbit 271_09, case of limb observation with slit approximately parallel to the limb. (top) Image at 1.27 mm. (bottom) Image of the same cube but at 1.74 mm, same scale. The isolines are the tangent altitude. (c) Orbit 44_01, case of nadir mode. Several spectra in the 1.27 mm range are shown for different points of the image. Different colors of spectra coincide with the same color of area in the red-green-blue (RGB) image. Positions of the R, G, and B bands are shown on the spectra. The area of O2 emission appears white in this composite. The shift to shorter wavelengths of the nightglow maximum emission compared to the maximum of the thermal emission from below the clouds can be seen (compare for example the green and magenta spectra). The small peak in the spectra at about 1.31 mm is a thermal emission in another atmospheric window.
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Table 1. Summary of the Selected Dataa Volume Emission Rate
Limb Session
Start Time
Latitude Range (deg)
VI0076_18b VI0271_09 VI0275_06 VI0317_06b VI0317_07b VI0317_08b VI0320_03b VI0320_04b VI0321_03b VI0321_04b VI0321_05b VI0322_06b VI0322_07b VI0322_08b VI0323_07b VI0323_08b VI0324_06b VI0324_07b VI0324_08b VI0327_05b VI0327_06b VI0327_07 VI0330_04b VI0330_05b VI0330_06 VI0364_07 VI0364_08 VI0364_09 VI0371_10b VI0377_11b VI0383_12 VI0443_06b VI0449_10b VI0457_08 VI0460_08 VI0461_07 VI0499_11b VI0507_11 VI0507_12b VI0516_05b VI0517_05b VI0519_05b VI0565_18b VI0565_19b VI0565_20b VI0600_00 VI0600_02b VI0602_00 VI0602_02b VI0623_00 VI0623_01b VI0718_03 VI0724_02 VI0733_02 VI0742_02 VI0754_01 VI0754_02 VI0764_01 VI0764_02 VI0766_01 VI0766_02 VI0792_01 VI0792_02 VI0796_02 VI0796_03 VI0802_00 VI0802_01 VI0802_02 VI0806_01 VI0806_02
2006-07-06T01:34:11.885 2007-01-17T07:36:25.808 2007-01-21T07:33:11.848 2007-03-04T06:12:02.447 2007-03-04T06:26:02.459 2007-03-04T06:40:02.483 2007-03-07T06:09:36.409 2007-03-07T06:23:36.415 2007-03-08T06:08:48.416 2007-03-08T06:22:48.431 2007-03-08T06:36:48.443 2007-03-09T06:07:58.525 2007-03-09T06:21:58.413 2007-03-09T06:35:58.536 2007-03-10T06:21:10.344 2007-03-10T06:35:10.475 2007-03-11T06:06:22.483 2007-03-11T06:20:22.504 2007-03-11T06:34:22.594 2007-03-14T06:03:57.504 2007-03-14T06:17:56.492 2007-03-14T06:31:57.442 2007-03-17T06:01:32.528 2007-03-17T06:15:32.429 2007-03-17T06:29:32.559 2007-04-20T05:36:38.960 2007-04-20T05:50:45.877 2007-04-20T06:04:52.891 2007-04-27T05:42:00.927 2007-05-03T05:37:23.199 2007-05-09T05:51:47.236 2007-07-08T07:36:05.904 2007-07-14T07:38:35.961 2007-07-22T07:14:22.048 2007-07-25T06:58:28.022 2007-07-26T06:53:09.967 2007-09-02T03:13:58.502 2007-09-10T02:34:10.606 2007-09-10T03:07:10.565 2007-09-19T02:04:08.786 2007-09-20T01:58:48.667 2007-09-22T01:53:38.770 2007-11-07T00:50:16.235 2007-11-07T01:12:15.305 2007-11-07T01:26:16.217 2007-12-12T02:07:05.806 2007-12-12T02:45:05.782 2007-12-14T02:10:56.781 2007-12-14T02:48:55.817 2008-01-04T02:56:19.960 2008-01-04T03:14:19.984 2008-04-08T04:36:17.693 2008-04-14T04:30:38.081 2008-04-23T04:06:09.731 2008-05-02T04:20:36.192 2008-05-14T02:54:24.287 2008-05-14T03:08:24.287 2008-05-24T03:02:37.862 2008-05-24T03:16:38.266 2008-05-26T03:05:07.274 2008-05-26T03:19:07.289 2008-06-21T03:37:10.343 2008-06-21T03:50:10.343 2008-06-25T03:42:10.367 2008-06-25T03:55:10.371 2008-07-01T03:32:44.307 2008-07-01T03:49:44.389 2008-07-01T04:03:44.411 2008-07-05T03:54:48.428 2008-07-05T04:08:48.562
30.0:65.0 15.0:35.0 15.0:40.0 15.0:30.0 25.0:50.0 50.0:75.0 15.0:35.0 25.0:50.0 15.0:30.0 25.0:50.0 50.0:75.0 15.0:30.0 25.0:50.0 60.0:70.0 30.0:50.0 50.0:60.0 15.0:30.0 25.0:50.0 50.0:60.0 15.0:30.0 35.0:50.0 50.0:75.0 15.0:30.0 25.0:45.0 50.0:75.0 15.0:30.0 30.0:45.0 55.0:65.0 35.0:60.0 35.0:60.0 35.0:50.0 0.0:20.0 0.0:20.0 0.0:25.0 0.0:25.0 0.0:25.0 10.0:40.0 10.0:40.0 0.0:20.0 0.0:30.0 0.0:25.0 0.0:25.0 10.0:15.0 25.0:40.0 45.0:70.0 10.0:25.0 50.0:65.0 10.0:25.0 50.0:75.0 10.0:25.0 20.0:50.0 0.0:20.0 0.0:5.0 5.0:30.0 5.0:15.0 25.0:45.0 45.0:70.0 25.0:45.0 45.0:70.0 25.0:45.0 45.0:70.0 25.0:40.0 45.0:70.0 25.0:40.0 45.0:70.0 10.0:20.0 25.0:40.0 45.0:65.0 25.0:40.0 45.0:70.0
Local Time Range (h)
Peak From Limb (km)
21.8:23.9 23.8:2.4 0.3:2.9 23.8:0.7 23.7:0.4 0.03:7.1 0.18:1.1 0:0.8 0.31:1.2 0.1:0.9 0.46:7.4 0.43:1.3 0.21:0.97 0.57:7.5 0.32:1.08 0.68:7.6 0.67:1.5 0.43:1.2 0.78:7.7 1.05:1.8 0.77:1.5 1.11:7.99 1.4:2.1 1.1:1.8 1.4:8.26 21.4:21.9 21.6:22.3 17.8:21.8 21.9:23.5 22.6:0.16 18.03:0.8 0.9:2.86 1.6:3.5 19.2:21.0 19.5:21.4 19.6:21.5 23.9:3.1 0.8:3.3 23.6:1.7 0.9:3.0 1.0:3.1 1.3:3.4 2.6:3.5 2.0:2.9 2.2:5.1 22.2:23.5 20.2:23.7 22.4:23.7 20.4:23.9 0.8:2.2 1.4:2.4 22.4:0.3 22.4:0.8 0.7:2.2 1.3:2.9 22.0:23.0 22.2:0.7 23.1:0.1 23.2:1.9 23.3:0.3 23.5:2.1 2.3:3.1 2.3:4.7 2.8:3.6 2.7:5.2 20.1:20.8 20.5:21.3 18.9:21.3 20.9:21.8 19.5:21.8
95.6 95.7 94.5 94.3 94.2 95.2 93.0 94.8 92.3 96.0 98.2 92.7 96.4 98.0 96.5 97.0 96.7 96.8 98.0 96.3 93.3 93.2 92.7 93.3 93.8 97.3 97.0 98.0 97.2 97.0 94.0 94.8 92.5 93.0 93.2 93.8 98.8 98.5 94.8 94.7 94.8 95.4 100.0 97.3 93.0 96.3 100.3 97.0 93.2 93.3 95.0 93.5 94.5 97.6 94.5 93.5 93.0 93.0 91.6 93.8 92.8 98.0 93.5 91.3 91.2 92.5 94.7 98.8 96.3 95.6
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Peak (km)
FWHM (km)
Peak Intensity (MR/km)
Total VER (MR)
96.7 97.3 97.3 96.0 96.6 97.4 95.3 96.2 95.3 98.3 101.4 96.3 99.4 101.0 99.8 101.0 98.7 99.6 101.0 101.0 95.0 94.6 95.7 95.0 95.4 99.7 99.3 98.5 99.6 99.6 95.3 98.3 95.8 96.6 96.4 97.0 100.2 100.3 98.3 96.8 97.0 98.2 103.0 98.7 95.0 99.3 101.3 101.7 94.4 96.7 97.3 98.0 98.5 99.8 97.5 95.8 95.0 95.8 93.6 96.3 94.2 100.7 95.0 94.7 94.0 95.0 96.3 99.8 98.3 96.8
3.9 5.7 6.3 9.7 9.0 5.8 8.5 6.2 9.3 8.7 6.4 9.3 6.6 5.0 6.3 7.0 9.3 7.8 8.0 11.0 7.3 6.4 9.7 7.0 5.6 7.7 7.0 3.5 5.6 7.4 6.7 9.8 10.3 10.0 10.4 9.8 6.5 7.2 10.3 9.7 9.0 8.0 9.0 6.0 7.4 8.0 3.7 11.0 5.4 9.3 7.3 11.0 12.5 8.4 10.0 7.3 6.5 8.8 8.2 7.0 6.8 8.3 7.5 8.7 7.6 9.0 6.0 4.3 9.3 4.6
0.06 0.14 0.06 0.12 0.05 0.04 0.03 0.04 0.03 0.06 0.13 0.03 0.03 0.09 0.08 0.14 0.10 0.08 0.08 0.12 0.03 0.02 0.02 0.03 0.02 0.09 0.11 0.10 0.11 0.10 0.04 0.11 0.08 0.01 0.01 0.02 0.10 0.11 0.11 0.10 0.15 0.12 0.09 0.04 0.01 0.08 0.14 0.06 0.02 0.03 0.06 0.05 0.10 0.14 0.07 0.06 0.01 0.03 0.02 0.03 0.02 0.07 0.03 0.01 0.01 0.06 0.05 0.16 0.10 0.04
0.27 0.91 0.43 1.10 0.49 0.26 0.31 0.26 0.29 0.45 0.85 0.25 0.24 0.47 0.64 1.01 0.96 0.59 0.58 1.19 0.28 0.10 0.23 0.21 0.11 0.73 0.77 0.48 0.78 0.81 0.29 1.03 0.90 0.12 0.13 0.17 0.71 0.87 1.16 1.03 1.30 0.99 0.86 0.27 0.10 0.74 0.69 0.68 0.15 0.32 0.44 0.58 1.32 1.28 0.73 0.47 0.11 0.26 0.17 0.24 0.11 0.67 0.25 0.10 0.09 0.55 0.37 0.74 0.98 0.26
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Table 1. (continued) Volume Emission Rate Local Time Range (h)
Peak From Limb (km)
Limb Session
Start Time
VI0843_03 VI0847_03 VI0320_05 VI0321_04 VI0323_06
2008-08-11T05:19:18.518 2008-08-15T05:28:28.552 2007-03-07T06:37:36.350 2007-03-08T06:22:48.431 2007-03-10T06:07:10.379
25.0:45.0 25.0:45.0 60.0:70.0 35.0:45.0 20.0:30.0
1.3:2.0 1.8:2.4 0.4:7.3 0.1:0.9 0.5:1.4
Nadir Session
Start Time
Latitude Range (deg)
Local Time Range (h)
Exposure Time (s)
VI0044_01 VI0093_00 VI0093_01 VI0093_02 VI0264_04 VI0344_01 VI0574_05 VI0586_01
2006-06-03T15:06:56.666 2006-07-22T15:00:09.328 2006-07-22T15:28:09.283 2006-07-22T15:56:09.355 2007-01-09T20:36:56.857 2007-03-30T23:32:32.734 2007-11-15T18:42:55.387 2007-11-27T16:53:15.434
79.4:9.9 59.1:2.5 60.9:7.3 45.2:11.3 78.6:10.6 66.2:9.3 40.1:15.7 62.3:8.6
19.4:2.7 1.3:4.7 20.9:1.5 19.3:22.8 19.6:3.1 19.6:23.7 21.5:0.4 21.6:2.6
3.3 3.3 3.3 3.3 3.3 18.0 8.0 8.0
93.5 94.3
Peak (km) 95.3 97.0 Limb profile Limb profile Limb profile
FWHM (km)
Peak Intensity (MR/km)
Latitude Range (deg)
6.3 7.0 with double with double with double
0.03 0.06 peak peak peak
Total VER (MR) 0.23 0.45
a List of the selected data in limb and nadir mode and some measured parameters. The session name is shown. The first two letters of the session name (VI) are fixed and they mean ‘‘Virtis Infrared’’; the four numbers after VI are the orbit number; the two suffix numbers after the underscore are the session number, which is the progressive observation number within the considered orbit. The start time of the acquisition is shown. Each limb observation lasts typically 30 – 40 min. The range of latitude spanned by the considered observation is given. The local time spanned range by the considered observation is given. The observed peak altitude for the limb profile, averaged over the latitude range of the session, is given (only limb mode). The parameters from the retrieved volume emission rate are the peak altitude, the FWHM, the peak intensity, and the integrated volume emission rate along the perpendicular direction. All the parameters reported are an average of all the grid point values within the range of latitude of each session. For the peak altitudes and the FWHM, an accuracy of ±1 km can be assumed. For the peak intensity and VER, an accuracy of ±10% can be assumed. The exposure time for all the limb observations is 8.0 s. b Used to infer the ratio 1.27 mm/1.58 mm. The ratio is calculated from a portion of the session where the intensity is brighter and not from the full session. For nadir data, the exposure time is 3.3, 8.0, or 18.0 s.
Figure 2b. Figure 2 includes limb profiles from the sessions shown in Figures 1a and 1b (sessions 76_18 and 271_09). [13] In nadir mode observations, the thermal emission cannot be neglected so it must be removed from the observed radiance to quantify the oxygen nightglow emission. Our first step in determining the thermal emission was to construct a false color composite red-green-blue (RGB) image using data from the 1.285 (red), 1.275 (green), and 1.265 (blue) mm wavelengths, as shown in Figure 1c. Regions dominated by nightglow emission appear blue or white. Those dominated by thermal emission are pink or red. To quantitatively remove the thermal contribution at 1.27 mm, we have scaled the 1.18-mm radiance by an empirical factor ranging from 0.27 to 0.38. This empirical scaling factor was determined by selecting spectra with a small intensity at 1.265 mm and a high correlation between the radiances at 1.18 and 1.285 mm (reddish areas without O2 emission on the RGB images). Variations in the scaling factor across the disk are due mostly to topography but there is a weak dependence on cloud structure. [14] The scaling factor to remove the thermal contribution at 1.27 mm also has been estimated using radiative transfer model calculations. Synthetic spectra were calculated using a surface temperature of 735 K and the typical cloud structures found at middle to low latitudes [Zasova et al., 2007]. The total vertical cloud opacity at 1.18 mm was varied from 20 to 50 by varying the opacity of each layer without altering the structure of the clouds, which were assumed to be composed of mode 1 and mode 2 particles with scale height of 4 km. The calculations used the ARS code [Ignatiev et al., 2005], which is based on DISORT
[Stamnes et al., 1988], with 48 streams and the HITEMP database [Pollack et al., 1993] for CO2 absorption. For the synthetic spectra, the ratio of the integrated emission in the 1.27- and 1.18-mm windows, r = I(1.27)/I(1.18), is in the range from 0.28 to 0.34, consistent with the empirical ratio from 0.27 to 0.38. The ratio depends only weakly on the total vertical cloud opacity because the single scattering albedo of the particles is very high (0.99999 at 1.18 mm). The single scattering albedo in the 1.74-mm window is smaller, 0.99978, so it does not work as well for the thermal correction. The ratio is most sensitive to the surface temperature. If the surface temperature is changed by 15 K, which is equivalent to a 2 km change in surface altitude, then the ratio changes by 0.03 for a cloud opacity of 30. [15] Vertical profiles of the volume emission rate are retrieved from the observed limb profiles using an onion peeling technique [Sharma et al., 1988]. Figure 3 shows the vertical profile of the volume emission rate of the (0 – 0) oxygen nightglow retrieved for orbit 317_06. The emission rate is expressed in MR/km (or 107 photons cm3). The technique assumes the atmosphere is spherically symmetric, including the airglow emission layer. Nadir-view images, however, indicate the nightglow emission is spatially inhomogeneous. The effective consequence of our assumption of spherical symmetry is to distribute the nightglow emission observed at higher altitudes over a larger region than may be true in reality. As a result, the retrieved emission at lower altitudes is underestimated. Clear evidence for this is found when the retrieved intensity below the peak altitude becomes negative. For several orbits where the spherical symmetry assumption led to negative values for the retrieved nightglow emission below 90 km, we imposed
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Figure 2. Limb profiles distribution of the O2 (0 –0) nightglow at 1.27 mm. (a) Limb profiles of radiance at 1.18 and 1.74 mm, shown in comparison with the O2 (0 – 0) limb profile at 1.27 mm. The radiance at 1.18 and 1.74 mm is due to the scattered thermal emission from the deep atmosphere scattered to VIRTIS-M by the upper haze particles. Both radiances decay rapidly in the transition region of the upper boundary of the aerosol at about 90 km altitude. The thermal contribution at 1.27 mm should have a similar behavior; therefore, in the typical altitude range where the nightglow emitter resides (90 – 105 km), an almost pure nightglow emission is observed in limb view. (b) The first three numbers of the curve names are the number of the orbit, while the given range after the orbit number is the latitude range where the average was taken. Data from all local times spanned in the session were included in the average. One can see that the peak altitude emission and its brightness significantly vary from one orbit to the next one and even in the same orbit. From Figure 2, the profile seems wider in altitude at lower latitudes, and a double peak is often present with a peak at 96 and 103 km.
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Figure 3. Vertical profiles of the O2 (0 – 0) retrieved volume emission rate. Black curves are volume emission rate in O2 (0 – 0) band at 1.27 mm, for orbit 317_06 in different latitude ranges; red curve is normalized radiance at 1.18 mm, which describes the upper boundary for the lower atmosphere thermal emission scattered into the VIRTIS-M field of view limb view. The normalized limb profile of thermal emission at 1.18 mm coincides well with the retrieved emission at 1.27 mm below 90 km, which is also thermal emission from the lower atmosphere. set of limb observations from different orbits over a period of about 2 weeks (in contour plot style). [17] These results reveal several characteristics of the O2(a-X) nightglow emission. As was previously known, the intensity of the emission is highly variable. Figure 2b shows the maximum observed peak limb intensity varying from about 15 to 80 MR, and the minimum and maximum observed peak limb intensities in the 72 sessions studied are 4 and 100 MR, respectively. Figure 2b also demonstrates that the observed altitude of peak brightness can vary. This
the additional constraint that the retrieved nightglow emission must be nonnegative. From these test cases, the upper limit on the geographic size of the emitting area was 500 – 1500 km, consistent with nadir images. When the retrieval is satisfactory, the shape of the vertical profile of the 1.27-mm radiance below 90 km coincides with that of the 1.18 mmradiance, both providing the profile of thermal radiance from the deep atmosphere as scattered in the upper haze layer.
5. Vertical Distribution of the O2(a-X) Nightglow Emission at 1.27 mm [16] In Figure 4 we show observations from a single orbit of the (0 – 0) limb profiles averaged over 5° latitude intervals which were also averaged over the full local time span at each latitude, about 10– 20 min; Figure 5 shows a selected
Figure 4. Limb profiles distribution of the O2 (0 – 0) nightglow in limb mode. The intensity in limb view for orbit 317 at 35 –73°N, session 7 and 8, is shown. Different peak altitudes are observed at different latitudes with a regular trend of correlation within this orbit. Also, the full width at half maximum (FWHM) appears wider toward the equator.
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Figure 5. Limb profiles distribution of the O2 (0– 0) nightglow at 1.27 mm versus latitude. The intensity is integrated over all the local times at each latitude. The color bar is the intensity along the line of sight. The orbit and session number are given in the titles. The intensity can differ significantly from one limb observation to another, as, for example, in the case of Figures 5a and 5b. The altitude can be almost constant, like in Figure 5f, and moderately constant like in Figure 5c, or more frequently a slope is observed like in Figures 5d and 5e. Some small ‘‘wave-like’’ features, as in Figure 5b at 27°N, are real and may possibly be attributed to gravity waves. is new information that cannot be obtained from nadir observations. In Figure 2b the observed peak altitude of the limb profile ranges from 96 to 103 km (for the higher peak of the double-peak profile), and the minimum and maximum altitudes for the observed peak of the nightglow limb profile in the 72 sessions studied are 90 and 102 km
(for the single-peak profiles), respectively. Further study is required to assess whether the variations in the observed peak altitude may explain the temperature variations derived from ground-based observations of the emission [Bailey et al., 2008]. Finally, significant variations have been detected in the shape of the observed limb profile (Figure 2b). Most
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observed limb profiles have a single peak, as has been reported previously [Drossart et al., 2007a; Ge´rard et al., 2008], but some observed limb profiles have a double peak. This has not been reported previously and is discussed later in this paper. [18] A view of the longer-term average characteristics of the (0 – 0) limb observations can be derived by using all of the limb data available in the VIRTIS data set to investigate the statistical characteristics of the retrieved vertical emission rate. The observed limb profile for each session, averaged over all local times (Table 1) was determined, assigned to a point on a 5° latitude range, and inverted to retrieve the vertical profile. Each retrieved vertical profile was then analyzed to derive the retrieved peak altitude, the peak altitude brightness, the FWHM of the retrieved vertical profile, and the total (column integrated) vertical emission rate. The mean values for each of these parameters for each latitude range are calculated and shown in Figure 6. Figure 6a shows the number of available limb observations for each latitude range. In Figure 6b, we show the meridional distribution of the mean peak altitude brightness. The abscissa is latitude, as for the other plots of Figure 6, and the ordinate is the intensity of the vertical emission rate at the retrieved peak altitude. The mean peak emission rate ranges from about 0.03 to 0.11 MR/km. Figure 6c shows the meridional distribution of the mean values of the altitudes at which the integrated line of sight (as observed, red curve) and vertical emission (as retrieved, blue curve) rates are maximum. For an ideally thin emitting layer, the observed peak altitude would coincide with the retrieved peak altitude. The observed peak altitude is systematically lower than the retrieved peak altitude by 2 – 3 km. Most of this bias is due to the geometrical effect of a not ideally thin emitting layer. In Figure 6d is shown the meridional distribution of the mean FWHM of the retrieved vertical profile of the oxygen nightglow. Figure 6e shows the total (column integrated) vertical emission rate, which, in turn, is equivalent to the intensity that would be seen in nadir view for pure oxygen nightglow emission. Finally, in Figure 6f, isolines of the mean total (column integrated) vertical emission rates are shown as a function of latitude and the retrieved peak altitude. The uncertainty in the peak altitudes and the FWHM is ±1 km, while that for the emission rate is ±10%. In all cases, however, the observed fluctuations are dominated by temporal variability in the nightglow emission. For this reason the error bar in the plots reports the standard deviation of this fluctuation. [19] While the mean values discussed above represent the general morphology, there are occasional but significant deviations. For example in many cases the peak intensity decreases with increasing latitude, but sometimes it is observed an inverse behavior, as the one for example shown in Figure 2b, sessions 321_04 and 321_05. In some orbits of Figure 5, the maximum of the nightglow intensity is at subequatorial latitudes but there are secondary local maxima at higher latitudes. For example, session 324_06 (Figure 5d) has one maximum near 30°N and orbit 330 (Figures 5e and 5f) has several maxima at 25 –45°N. Figure 5 also suggests an episodic variation in the maximum nightglow intensity with time for the region that was observed. The maximum intensity decreased by a factor of 4 over four days from session 317_06 (Figure 5a) to session 321_03 (Figure 5b)
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with almost the same local time; increased over the next three days, so that the observed maximum intensity in session 324_06 (Figure 5d) is similar to that in session 317_06; and then decreased again in session 330_04 (Figure 5e) to values similar to those observed in session 323_06 (Figure 5c). Some dynamical factors, such as breaking gravity waves at a local scale and tides at larger scales may disrupt the general picture, so long-term averages may reveal different dynamical information than the single orbit snapshots. [20] The spatial distribution of the nightglow intensity derived from limb observations (Figure 6) is consistent with that found from the previous analysis of VIRTIS’ nadir observations [Ge´rard et al., 2008] and our analysis of a more complete set of nadir observations, as discussed later. In general, there is a negative correlation between the intensity of the emission and latitude. As shown in Figure 6b, the most intense emission observed is found at low latitudes. The peak altitude and the FWHM of the nightglow’s vertical profile also show regular trends with latitude. As shown in Figure 6c, the mean retrieved peak altitude is highest at 45– 50°N with a value of about 99 km, and appears to be relatively constant equatorward with a value of 97– 98 km, while it decreases poleward with a relative minimum of 95 km at 65– 70°N. The vertical FWHM of the emission shows a clearer trend, increasing from high to low latitudes with the largest FWHM, 11 km, at 5– 10°N and the smallest, 6 km, at 60– 65°N (see Figure 6d). The decrease of the observed peak altitude and its larger deviation (to lower altitudes) from the retrieved peak altitude toward the equator, see Figure 6c (left), can be explained by the increase of the vertical FWHM in this region. The thicker the emitting layer, the larger the difference between the observed and retrieved peak altitudes. A similar trend is found in Figure 4, where the observed peak altitude of the nested profiles increases with latitude, while FWHM and intensity decrease. The decrease in the peak altitude poleward of 45°N (Figure 6c) may indicate a different dynamical regime exists at the altitude of the emitting layer. This may be related to the influence of the polar vortex. [21] Regions of intense nightglow where the peak altitude is approximately constant are likely to be regions of convergence for the subsolar to antisolar flow with strong downwelling. In terms of pressure, a constant peak altitude is consistent with the conditions to have very low horizontal pressure gradient forces which will suppress horizontal flow and give predominantly vertical flow. This is illustrated by the distribution of the total (column integrated) vertical emission rates derived from our limb observations (Figure 6e) and the distribution of peak altitudes (Figure 6c). The former increases by a factor of 6 from about 0.2 MR at 60–70°N to about 1.2 MR near the equator. The latter is relatively constant equatorward of 45°N. The overall picture derived from our limb observations is consistent with that derived from our own and previous analyses of nadir observations. In addition, we have been able to determine that the enhanced total (column integrated) vertical emission rate near the equator is due to both a more intense vertical emission rate at the retrieved peak altitude (Figure 6b) and a larger vertical FWHM for the emitting layer (Figure 6d). The former contributes a factor of three to the equatorial enhancement and the latter a factor of two. Finally, as shown in Figure 6f, the most intense O2 nightglow emission is
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Figure 6. Mean vertical profile distribution of the O2 (0 – 0) nightglow at 1.27 mm. (a) Number of available sessions in the latitude range is shown to indicate the number of samples over which the average is calculated for each latitude range. (b) Mean emission rate at peak altitude as a function of latitude. The error bar is the standard deviation. The maximum is found in the equatorial region, and the intensity decreases poleward. (c) Mean peak altitude of the volume emission rate (blue curve) and in limb profile (red curve) as a function of latitude. The error bar is the standard deviation. Figure 6 indicates the peak altitude of the volume emission is relatively constant and tends to decrease poleward of 45– 50°N. The weak decrease of the peak altitude of the limb profile (red curve) is likely related to the wider FWHM at low latitudes. (d) Mean FWHM of the vertical profile as a function of latitude. The error bar is the standard deviation. The results show the FWHM is largest at low latitudes and decreases poleward with some relative peak at 45– 50° and 70°N. (e) Mean total integrated vertical emission rate as a function of latitude. The error bar is the standard deviation. The maximum is found in the equatorial region, and the intensity decreases poleward. The slope is less pronounced than in Figure 6b owing to the amplification factor of Figure 6d. (f) Same as Figure 6e but as a function of latitude and peak altitude of the vertical profile. An uncertainty of ±10% in emission can be assumed. The intensity is highest at low latitudes when the peak altitude is higher.
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Figure 7. Population distribution of the oxygen nightglow parameters. (a) Distribution of the peak altitude emission rate intensity, same parameter as Figure 6b. It is remarkable the large variability of the intensity which is spread over a wide range of emission. The global mean value of the peak emission rate intensity is 0.064 ± 0.050 MR/km. (b) Distribution of the peak altitude of the vertical profile. Variability is less pronounced than in Figure 7a, and the global mean value of the peak altitude of emission is 97.4 ± 2.5 km. (c) Distribution of the FWHM of the vertical profile. The global mean value of FWHM is 7.6 ± 2.2 km. (d) Distribution of the total integrated vertical emission rate. Also, in this case the large variability is noticeable. The global mean value of the integrated emission rate is 0.52 ± 0.40 MR. observed at near equatorial latitudes when the retrieved peak altitude is at least 98 km. These are the most favorable conditions for radiative deexcitation as opposed to collisional quenching. [22] The statistical distributions of the parameters are shown in Figure 7. In average, the peak vertical emission rate intensity (peak altitude brightness) is 0.064 ± 0.050 MR/km (Figure 7a); the retrieved peak altitude is 97.4 ± 2.5 km and the observed peak altitude is 95.0 ± 2.4 km (Figure 7b); the vertical FWHM is 7.6 ± 2.2 km (Figure 7c); the total (column integrated) vertical emission rate is 0.52 ± 0.40 MR (Figure 7d). The observed and retrieved peak altitudes vary by ±2.5% across the range of latitudes observed while the total (column integrated) vertical emission rate varies by more than ±75%. A factor of 50 has been inferred from the limb to the perpendicular direction views of the intensity of the emitting layer in average.
6. Detection of Double-Peaked Nightglow [23] In many cases two emission peaks are found in the observed limb profiles. Some examples are shown in Figures 2 and 8. In Figure 8a the measured limb profiles are shown on the left and the corresponding limb images at 1.27 mm on the right. The separation of the two layers is
clear. On the occasions when a double peak is detected, the lower and higher peaks occur typically at 96– 98 and 103– 105 km, respectively, see Figure 8a and Figure 8b (left) where the retrieved vertical profiles are shown. Double peaks are more likely to occur at high latitudes, 45– 70°N, although there are also occasional examples at lower latitudes. The higher peak is often the more intense of the two. [24] The origin of the double peak is not yet understood. However, some suggestions can be provided. Given the uneven global distribution of the nightglow, one may think that the higher and lower peaks are the results of probing two separate emitting layers, each at a different altitude level. If so, the double peak would be a consequence of the viewing geometry. It is also possible the viewing geometry could produce an apparent double peak if both emitting regions are at the same radial distance from Venus but with a gap between them where there is less nightglow emission. Supposing this to be true for the case shown in Figure 8a, for the session 320_05, let’s assume the higher peak as coming from a single-peaked layer located in a region at the LOS tangent altitude. The lower peak might be explained then with a local maximum of another emitting region placed off by 3° (about 310 km) from the former and at the same radial distance. However, this region of local maximum, estimated from the black curve of Figure 8a to
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Figure 8. Profiles of the oxygen nightglow with double peak. (a) Comparison of different limb profiles for the (left) O2 (0 – 0) nightglow at 1.27 mm and (right) the image view of the same observations. The green isolines in the images are the tangent altitude. The yellow isolines of the same images are latitudes, and they mark the latitudinal range where the limb profiles are taken. Color scale intensity is not the same for the images, in order to optimize the contrast. Double peaks are noticed in the limb profile on the left plots, and on the right, spatially separated layers can be seen in the acquired images. (b) Vertical emission rate retrieved for the profiles shown in Figure 8a. The double peak is observed also for the retrieved profiles and it results even more enhanced with relative peaks very well separated vertically.
be about 2.5 times brighter in total VER than the region responsible of the higher peak, should develop in a long strip precisely at this same distance for more than 20° latitude, corresponding to about 2000 km of spatial displacement (see Figure 8a, top right). Therefore, although these combinations can reproduce a double-peaked limb profile in some cases, this scenario appears to be very unlikely to explain the two distinct emitting layers of Figure 8 and several other similar observed cases not reported here. [25] A third possible explanation, which is somewhat simpler and appears to be more likely than the other two, involves vertically propagating gravity waves. Gravity waves can modulate the conditions in the atmosphere, thereby modifying density, temperature, atomic oxygen abundance, and hence the collisional rate of recombining oxygen atoms and collisional quenching of the O2(a) state. As a matter of fact, wave structures at Venus were observed by the Pioneer Venus Orbiter star trackers [Bougher and
Borucki, 1994] in the same range of middle and high latitudes where double peaks are often found. Finally, it must be said that the retrieved atomic oxygen number density vertical profile is often double peaked even when the observed limb profile is single peaked [Ge´rard et al., 2009].
7. O2 Emission Rate From Nadir Observations [26] Figure 9 shows a nadir-looking view of the nightglow. The emission rate has been corrected for thermal emission in the way described above, while emergence angle and cloud backscattering are corrected as was done by Crisp et al. [1996]. Various structures with clearly defined boundaries are visible in Figures 9a– 9d that denote the complex interplay of dynamical and, possibly, chemical factors in the formation of the nightglow. Notwithstanding the significant variability observed, it seems that the distri-
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Figure 9
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bution of the nightglow over the planet is not entirely random but shows structured recurrent features. The emission patches may appear elongated, as seen in Figures 9a– 9c, in the form of jets over extensions of 1700 to 2300 km, or as wavy features as in Figure 9d. [27] We have searched for a connection between the nightglow and the local temperature. For this purpose, we have retrieved the average thermal brightness (TB) from the radiance in the 4.23 – 4.28 mm spectral interval which probes, according to radiative transfer models [Grassi et al., 2008], the 90– 93 km altitude range. Our approach is somewhat simpler than that used in the formal retrieval in that work, but the TB provides reliable relative values. In Figures 9e and 9f, intense nightglow seems to be correlated with elevated temperatures. Other authors have reported similar conclusions [Ohtsuki et al., 2005, 2008; Bailey et al., 2008; Bertaux et al., 2007], and it has been suggested that the intense nightglow may be caused by strong downwelling and resulting compressional heating. Conversely, the dim nightglow regions, if affected by upwelling, may be involved in expansion cooling with consequently lower temperatures than the surrounding environment. The same effects may also be a consequence of the atmospheric temperature structure modulation induced by vertically propagating gravity waves. The circular dark region of faint nightglow has on average a local TB of 182 K, lower by about 3 K than the surroundings with a TB of 185 K. The TB may differ by a few K from the formally retrieved actual absolute atmospheric temperatures, but their relative accuracy is estimated to be better than 1 K, so these temperature differences are considered significant. The bright region at the top (Figure 9f) where TB is on average about 189 K, corresponds to the region in Figure 9e where the nightglow is brightest and therefore the downwelling is a maximum in this region. [28] We have also performed a long-term study of the nightglow global distribution. For this purpose, a series of observations made with exposure times of at least 0.1 s and emergence angles smaller than 80° have been selected from 880 orbits. The radiances are subsequently corrected for thermal radiation (a constant factor of 0.3 was used in this case for scaling the 1.18 mm radiance), emission angle and cloud backscattering as previously discussed. The resulting image is shown in Figure 10a. The brightest region has a total (column integrated) vertical emission rate of 1.2 MR, consistent with the maximum derived from limb observations (Figure 6e). The center of the brightest region is about 5°S of the AS point, the proximity to which is an expected outcome of the subsolar to antisolar circulation. Figure 10b
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shows the map of coverage for the data used in Figure 10a, where the color scale bar represents the log10 of counts.
8. Oxygen Band Ratio [29] The (0 – 1) band of the a-X system that occurs at 1.58 mm is weaker than the 1.27 mm band. It is the former, however, that was first observed in the terrestrial atmosphere in the evening twilight ground experiment by Vallance Jones and Harrison [1958]. The atmospheric transmission at 1.27 mm is significantly lower, making it difficult for photons at this wavelength to reach the Earth’s surface. Some years would have to pass before the in-flight detection of the (0 –0) band by Noxon and Vallance Jones [1962], and some more for Lowe [1969] to resolve it in a spectrum recorded from the ground. The nondetection of the (0– 0) band in the pioneering work of Vallance Jones and Harrison [1958] was interpreted by the authors to mean a ratio of transition probabilities A00/A01 = 10, which is the value still in place in the current version of the HITRAN database [Rothman et al., 2005]. Observations of the (0 – 1) band in the terrestrial atmosphere continue, driven by the interest in investigating the oxygen chemistry of the mesosphere and lower thermosphere. [30] There have been a few subsequent estimates of the ratio, that have come either from laboratory experiments, e.g., A00/A01 = 46.0 ± 0.7 [Findlay, 1969], 49.5 ± 4 [Becker et al., 1971], 80 ± 20 [Haslett and Fehsenfeld, 1969], and 55 ± 10 [Thomas and Thrush, 1977]; or from airglow observations, some times in combination with photochemical modeling, e.g., A00/A01 = 52 ± 25 [Winick et al., 1985], and 80 ± 15 [Pick et al., 1971]. More recently, Kassi et al. [2005] have determined the line strengths of a few lines belonging to the strong Q branch in a high-resolution absorption spectrum experiment. By comparison to the respective lines in HITRAN, and after the correction of a numerical inconsistency in the conversion of absorption strengths to transition probabilities (confirmed by the authors), the work of Kassi et al. [2005] leads to A00/A01 = 55– 60. Recalling that on the basis of the Franck-Condon principle the predicted ratio is A00/A01 = 148 [Kassi et al., 2005] and that the principle is less likely to operate in longlived states, it is important to note the fundamental interest that the discrepancy between theory and experiments bears. [31] The VIRTIS instrument can resolve simultaneously the absolute radiances from both emission bands in the nighttime spectrum of Venus’ upper atmosphere. This was utilized by Piccioni et al. [2008] to infer A00/A01 = 63 ± 6. We now provide the ratio from a more extended data set and also how the ratio was determined.
Figure 9. Observations of the oxygen nightglow in nadir view geometry. O2 (0 –0) nightglow observations at 1.27 mm, after thermal, emergence angle, and cloud backscattering corrections: (a) orbit 264, session 04; (b) orbit 586, session 01; (c) orbit 574, session 05; (d) orbit 344, session 01. The green and yellow isolines are latitude and local time, respectively. (e) Orbit 93, mosaic of session 00, 01, and 02. The isolines from left to right are local time while the others from top to bottom are latitude. (f) Thermal brightness of the same observation as in Figure 9e, averaged from 4.23 to 4.28 mm which probes, according to radiative transfer calculation, the 0.1 mbar pressure level, corresponding to about 90– 93 km altitude. The green and red isolines are local time and latitude, respectively. The extension of the long stripes in Figure 9a is of the order of 1700 – 2300 km. The thermal brightness (TB) of the dark patch indicated by the arrow is about 182 K, which is 3 K lower than the surrounding. This is visibly correlated with the corresponding dark patch in Figure 9e of the oxygen nightglow. The diagonal dimension of the dark circular feature is about 800 km. The brighter region in the top middle of Figure 9e has the highest TB of the image, about 189 K, which is consistent with the maximum downwelling zone. 14 of 18
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bution of the thermal emission scattered by the clouds and haze, only observations with tangent altitude in the 90– 100 km range have been used. The continuum has been subtracted by linearly interpolating the observed radiances between 1.24 and 1.3 mm for the (0 – 0) band and 1.55 and 1.60 mm for the (0 –1) band. At the relevant temperatures, no contribution from the oxygen nightglow is expected outside these intervals. Subsequently we integrated the radiances for each transition to obtain the total emission rate. In Figure 11a the observed limb profiles from session 317_06 of the (0 – 0) and (0 – 1) bands at 1.27 and 1.58 mm, respectively, are shown as an example. In Figure 11b the ratios of the integrated radiances are shown. The mean radiance ratio integrated over the selected altitude range of 90– 100 km, and weighted by the inverse square of the uncertainty is 78 ± 10, which results in 63 ± 8 for the ratio of transition probabilities A00/A01. The uncertainty is largely due to the difficulty of evaluating the continuum of the 1.58-mm band so it is largest when the SNR in this spectral region is low.
9. Conclusions
Figure 10. Global map of the oxygen nightglow over 880 orbits in nadir geometry. (a) Mean global O2 (0 –0) nightglow at 1.27 mm, after thermal, emergence angle, and cloud backscattering corrections, averaged over 880 orbits. The map is constructed with all the available nadir observations in nighttime with an exposure time greater or equal than 0.3 s. The region of maximum emission, 1.2 MR, is near the antisolar (AS) point, a clear indication of the subsolar (SS) to AS circulation, but with a slight shift of the patch center toward the morning sector, to about 0015, and south, to 5°S. (b) Counts of the coverage map for the same data set as Figure 10a. Color scale bar is log10 (counts). The geometry of the orbit permits more frequent observations of the southern polar region owing to the longer persistence and farther distance of the apocenter which is located approximately over the south pole. However, it is shown that even in the less covered portion in the equatorial region, hundreds of observations are present. [32] For the analysis we have used portions of the 36 limb observations from Table 1 with relatively high intensity for the (0 – 0) band, of the order of 0.5 MR or greater, in the period March 2007 to January 2008. To avoid the contri-
[33] A study of Venusian molecular oxygen nightglow in the (0 –0) and (0 – 1) bands at 1.27 and 1.58 mm, respectively, from limb and nadir geometry has been conducted with the VIRTIS instrument on board Venus Express. Limb geometry data in the period from July 2006 to August 2008 as well as some nadir data, acquired in the period from 3 June 2006 to 27 November 2007, have been shown, described, analyzed, and compared with the thermal brightness at 4.23 to 4.28 mm, which probes the 0.1 mbar level of pressure, corresponding to about 90– 93 km altitude. All nadir data acquired in the period from 14 May 2006 to 14 September 2008, about 4 Venusian years, have been used to globally map the 1.27 mm emission averaged over this period of time (after correction for thermal emission, emergence angle and cloud backscattering). [34] The shape and the intensity of the 1.27 mm vertical profile are highly variable, however some regularities have been identified. The mean distribution of the total (column integrated) vertical emission rate inferred from limb observations is consistent with that derived from nadir observations over a 2-year period with the most intense nightglow emission found near the antisolar region. The retrieved peak altitude is relatively independent of latitude equatorward of 45– 50°N, with a mean value of 97.4 km while a different dynamical regime is identified poleward of this latitude. The vertical FWHM shows a significant increase with decreasing latitude and it reaches its maximum of 11 km in the equatorial region. The observed emission is most intense at low latitudes when the retrieved peak altitude is above 98 km. It has been shown that the intense total (column integrated) vertical emission rate that is observed near the antisolar point is due to both a more intense volume emission rate at the retrieved peak altitude and a thicker emitting layer. Limb profiles with two peaks have been observed. These could be related to vertically propagating gravity waves modulating the atmospheric structure. [35] Significant variability in the instantaneous local distribution of 1.27 mm emission rate in nadir mode is found, however features frequently appear very well struc-
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Figure 11. O2 (0 – 0) and (0 – 1) limb profiles and radiance ratio. (a) Intensity of limb profiles for total integrated emission in the (0 – 0) and (0 – 1) O2 transitions at 1.27 and 1.58 mm are shown for session 317_06. The intensity was averaged over the full range of latitude and local time of the session. The red scale shows the intensity value for the (0 – 1) transition, in MR. Their peak altitudes are in agreement. (b) O2 (0 –0) and (0 – 1) radiance ratio in limb mode at 1.27 and 1.58 mm is shown. The statistic takes in account all the data analyzed in the period March 2007 to January 2008 marked with a superscript "b" in Table 1. The ratio is
calculated in a selected bright region of the session only and not over the full spanned latitude and local time range of the session. This is why there is no direct correspondence from the ratio calculated from Figure 11a and the relevant point of the session in Figure 11b. The weighted mean value of the radiance ratio is 78 ± 10, which translates into 63 ± 8 in terms of the ratio of the transition probabilities A00/A01. tured, with very elongated bright stripes similar to jets and wavy type formations. A correlation of the nightglow intensity and the thermal brightness probed at the altitude of about 90– 93 km is often found in the instantaneous snapshot. Dark circular shape features occur frequently in the nightglow map and usually have a thermal brightness that is lower than the surrounding bright nightglow region. This suggests these dark regions may be places at which upwelling flow has inhibited the recombination of oxygen atoms and produced localized cooling. These dark features could also be produced by vertically propagating gravity waves. Finally, both the (0 – 0) and (0– 1) bands have been simultaneously measured in the VIRTIS spectra and a ratio of the transition probabilities A00/A01 = 63 ± 8 has been inferred.
[36] Acknowledgments. We gratefully acknowledge the work of the entire VIRTIS and Venus Express teams that allowed these data to be obtained and discussed. We wish to thank ASI, CNES, and the other national space agencies that have supported this research. Russian coauthors acknowledge Russian Foundation of Basic Research for grant RFFI 08-02-00850. Partial funding for this research was provided by the Australian Research Council. Helpful comments from two anonymous reviewers, D. Titov and R. Carlson, are gratefully acknowledged.
References Alexander, M. J. (1992), A mechanism for the Venus thermospheric superrotation, Geophys. Res. Lett., 19, 2207 – 2210, doi:10.1029/92GL02110. Allen, D. A., and J. W. Crawford (1984), Cloud structure on the dark side of Venus, Nature, 307, 222 – 224, doi:10.1038/307222a0. Bailey, J., V. S. Meadows, S. Chamberlain, and D. Crisp (2008), The temperature of the Venus mesosphere from O2(a1Dg) airglow observations, Icarus, 197, 247 – 259, doi:10.1016/j.icarus.2008.04.007. Becker, K. H., W. Groth, and U. Schurath (1971), The ratio of the FranckCondon factors q(0, 0)/q(0, 1) of the infrared atmospheric band system of
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Vallance Jones, A., and A. W. Harrison (1958), 1Dg X3S g O2 infrared emission band in the twilight airglow spectrum, J. Atmos. Terr. Phys., 13, 45 – 60, doi:10.1016/0021-9169(58)90024-2. Winick, J. R., R. H. Picard, R. D. Sharma, and R. M. Nadile (1985), Oxygen singlet Delta 1.58-mm (0 – 1) limb radiance in the upper stratosphere and lower mesosphere, J. Geophys. Res., 90, 9804 – 9814, doi:10.1029/JA090iA10p09804. Zasova, L., et al. (2007), Structure of the Venus atmosphere, Planet. Space Sci., 55, 1712 – 1728, doi:10.1016/j.pss.2007.01.011.
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P. Drossart, LESIA, Observatoire de Paris, UPMC, Universite´ ParisDiderot, CNRS, 5 place Jules Janssen, F-92195 Meudon, France. F. P. Mills and A. Garcı´a Mun˜oz, Research School of Physics and Engineering, Australian National University, Canberra, ACT 0200, Australia. A. Shakun and L. Zasova, Space Research Institute of Russian Academy of Sciences, Profsojuznaja 84/32, 117997 Moscow, Russia.
A. Cardesin-Moinelo, A. Migliorini, and G. Piccioni, INAF, Istituto di Astrofisica Spaziale e Fisica Cosmica, via del Fosso del Cavaliere 100, I-00133 Rome, Italy. (
[email protected])
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Limb observations of the ultraviolet nitric oxide nightglow with SPICAV on board Venus Express J.-C. Ge´rard,1 C. Cox,1 A. Saglam,1 J.-L. Bertaux,2 E. Villard,2 and C. Nehme´2 Received 17 January 2008; revised 15 April 2008; accepted 30 May 2008; published 20 August 2008.
[1] Limb observations of the spectrum of nightglow emission in the d (190–240 nm) and
g (225–270 nm) bands of nitric oxide have been made with the Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV) ultraviolet spectrometer on board Venus Express. These emissions arise from radiative recombination between O(3P) and N(4S) atoms that are produced on the dayside and recombine to form excited NO molecules on the nightside. No other emission feature has been identified. The mean altitude of the emission layer is located at 113 km, but it varies between 95 and 132 km. The mean brightness of the total NO emission at the limb is 32 kR, but it is highly variable with limb intensities as large as 440 kR observed at low latitude and values below 5 kR seen at northern midlatitudes. No systematic dependence of the brightness with latitude is observed, but the mean altitude of the emission maximum statistically drops with increasing latitude between 6° and 72°N. Typical observed limb profiles are compared with simulations based on a one-dimensional chemical-diffusive atmospheric model. From model fits to observed profiles, we find that the downward flux of N atoms at 130 km typically varies between 1 108 to 4 109 atoms cm2 s1. Comparisons of observed airglow topside scale heights with modeled profiles smoothed by the instrumental field of view indicate that the observations are compatible with a downward flow of O and N atoms by molecular and turbulent transport above the peak of emission. The K coefficient deduced from comparisons to limb profiles is less than that determined from the observations made with the Pioneer Venus UV spectrometer at low latitude during periods of high solar activity. Citation: Ge´rard, J.-C., C. Cox, A. Saglam, J.-L. Bertaux, E. Villard, and C. Nehme´ (2008), Limb observations of the ultraviolet nitric oxide nightglow with SPICAV on board Venus Express, J. Geophys. Res., 113, E00B03, doi:10.1029/2008JE003078.
or
1. Introduction [2] The presence of the delta and gamma bands of nitric oxide in the Venus nightglow was detected and identified by Feldman et al. [1979] using the ultraviolet spectrograph on board the International Ultraviolet Explorer (IUE). It was also observed by Stewart and Barth [1979] in spectra obtained with the ultraviolet spectrometer on board the Pioneer Venus Orbiter (PV-OUVS). The emission process is radiative recombination through reverse predissociation of nitrogen N(4S) and oxygen O(3P) atoms, yielding excited NO* molecules which emit the ultraviolet d and g bands: N þ O ! NO C2 P
ð1Þ
giving rise to NO C2 P ! NO X2 P þ d-bands;
1 Laboratoire de Physique Atmosphe´rique et Plane´taire, Universite´ de Lie`ge, Lie`ge, Belgium. 2 Service d’Ae´ronomie du CNRS, Universite´ Versailles Saint Quentin, Verrie`res-le-Buisson, France.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003078$09.00
NO C2 P ! NO A2 S; v0 ¼ 0 þ 1:22 mm;
followed by NO A2 S; v0 ¼ 0 ! NO X2 P þ g-bands
[3] As a consequence, the total emission rate of the NO bands is proportional to the rate of recombination of O and N atoms and thus depends on the nitrogen and oxygen densities. A detailed study of reaction (1) [Dalgarno et al., 1992] suggests a rate coefficient for the N + O recombination equal to 1.92 1017 (300/T)1/2 (1– 0.57/T1/2) cm3 s1. The N and O atoms are produced by dissociation of N2, CO2 and CO on the dayside. [4] Spin-scan images in the wavelength of the d (0,1) band at 198 nm obtained from Pioneer Venus near apoapsis by Stewart et al. [1980] indicated important day-to-day variations. Patches of enhanced intensity appeared to vary in intensity and location without correlation with solar activity. The location of the brightest spots ranged from 2130 to 0300 LT and 39°S to 60°N. This variability was attributed to temporal charges in the Venus thermospheric circulation on
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timescales of one Earth day or less. It appears to be caused by instabilities in the large-scale circulation, possibly as a result of wind shears near the terminator or time-varying wave drag from gravity waves. However, no quantitative modeling for this variability has been made so far. Stewart et al. [1980] also showed that, statistically, the emission is concentrated in a bright spot located near 0200 local solar time, just south of the equator. Following revision of the calibration of the instrument [Bougher et al., 1990], the emission rate of this region was estimated to 1.9 kR. From these observations, a general picture emerged where production of O and N atoms by solar EUV and possible solar particle impact occurs on the dayside, followed by global transport to the nightside, downward transport and radiative recombination. It was found that parameters such as the downward flux of O and N atoms and the strength of turbulent transport in a one-dimensional model could be adjusted to reproduce the average NO nightglow intensity. Since the attitude of the spacecraft was not known with sufficient accuracy, the determination of the altitude of the emission peak was indirect and based on the absorption of the NO emission by CO2. Ge´rard et al. [1981] observed that it lies close to 115 km. They showed that downward transport by molecular diffusion alone is not sufficient to reproduce the observations. They parameterized turbulent mixing using an eddy diffusion coefficient K 8 1012/n1/2 cm2 s1, where n is the total number density. This value was close to that deduced by von Zahn et al. [1979] from the analysis of the neutral composition measured near the morning terminator, which sets the altitude of the homopause close to 135 km. [5] Globally, the general picture of production of O and N atoms followed by transport to the nightside by the subsolar to antisolar circulation, downward turbulent mixing and radiative recombination appeared quantitatively consistent with the PV-OUVS observations. Further quantitative validation of this scheme was obtained by three-dimensional simulations using the Venus Thermospheric General Circulation Model (VTGCM) [Bougher et al., 1990]. Onedimensional chemical-transport calculations of odd nitrogen species on the dayside indicated that the daytime N average production is 1.3 1010 atoms cm2 s1 for solar maximum activity (F10.7 = 180 – 200) [Ge´rard et al., 1981]. Parameter adjustments in the VTGCM lead to an average production of about 9 109 atoms cm2 s1, approximately 30% of which are needed for transport to the nightside to reproduce the observed average nightglow brightness. The statistical location of the bright spot was also reasonably well reproduced by the 3-D model and implied zonal winds of about 50– 75 m s1 in the 115– 150 km region. However, it was found that model parameters could be adjusted to reproduce either the ultraviolet nightglow or the observed N density on the dayside but not both. Predictions for solar minimum conditions indicated a reduction by a factor of 3 from solar maximum. This is a direct consequence of the lower production of N and O atoms on the dayside and of the predicted reduced strength of the thermospheric circulation. The observed shift toward dawn of the statistical location of the airglow maximum, presumably as a consequence of the residual atmospheric superotation in the upper mesosphere, was reproduced by the VTGCM.
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[6] The NO d and g bands were also detected on Mars by Bertaux et al. [2005] using the SPICAM (Spectroscopy for Investigation of Characteristics of the Atmosphere of Mars) spectrometer on board the Mars Express spacecraft. The distribution and variability of the NO Mars airglow were recently described by Cox et al. [2008]. They found that the peak of the airglow layer is located between 55 and 92 km with lower values observed at higher latitude and larger values at low latitudes and midlatitudes. The limb brightness was also found to be very variable, ranging from less than 0.2 to 10.5 kR with the large values occurring at southern latitudes higher than 50°. [7] The NO airglow is thus a tracer of the thermospheric circulation and its distribution and intensity provide constraints on transport models in a region of the Venusian atmosphere difficult to probe with other in situ or remote sensing methods [Bougher et al., 2006]. In this study, we describe new observations of the NO nightglow collected in the northern hemisphere with the Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV) spectrometer on board the Venus Express spacecraft [Svedhem et al., 2007]. In particular, we investigate the distribution of the altitude and brightness of the airglow and its spatial variability. We examine possible relationships between the emission intensity and its brightness and a possible latitudinal dependence. We also compare limb profiles with results of a one-dimensional chemical-diffusive model solving the odd nitrogen and oxygen continuity equations and draw conclusions concerning the importance of transport by other processes.
2. SPICAV Observations [ 8 ] The observations described in this study were obtained with SPICAV, the ultraviolet and infrared spectrometer on board Venus Express. The instrument and its performances were described by Bertaux et al. [2007a]. The ultraviolet component covers the range extending from 118 nm to 320 nm and includes the NO d (C 2P - X 2P) and g (A 2S - X 2P) emission bands. For reasons of telemetry limitations but also because of the time needed to read all the lines of the CCD, only 5 adjacent zones of the CCD detector are usually read out. Each zone (or spatial bin) is made of 1, 2, 4, 8, 16 or 32 lines of the CCD following a preselected mode. Therefore each spatial bin covers a different region of the atmosphere separated by an angle ranging from 1.4 to 22.4 arcmin, depending on the spatial binning. In the case of the NO airglow limb scans reported here, each spatial bin includes 2, 4, 8, 16 or 32 adjacent pixel lines. These lines are seen through the large (500 mm) or the small (50 mm) slits, which provides a spectral resolution of 12 or 1.5 nm, respectively. The SPICAV CCD is read every second and therefore five spatial bins corresponding to five adjacent portions of the SPICAV field of view are recorded each second. In each individual spectrum non-uniform dark current and offset values are subtracted. Finally, the intensities are calibrated in kiloRayleighs (kR) using well-known hot stars spectra observed by SPICAV during the mission. [9] The VEX spacecraft moves along a quasi-polar eccentric orbit with a 24-hour period. The pericenter and apocenter are located at 250 km and 66,000 km,
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Figure 1. Variation of the altitude of the tangent point of the center of the field of view for SPICAV spatial bins 1 and 5 during tangential limb nightside observations of orbit 322. The altitudes of the other three spatial bins are within the gray shaded area. respectively, and the orbit precesses in local time at the rate of 1.6°/day. The precession of the orbital plane leads to a wide variety of configurations on the nightside as well as on the dayside. Several observation modes (‘‘science cases’’) may be selected including nadir observations, star pointing for stellar occultations by Venus’ atmosphere, fixed point tracking and limb observations [Titov et al., 2006]. A particular mode is well suited for high spatial resolution airglow observations. In this grazing (tangential) limb mode, the line of sight is at some angle to the orbital plane and moves in such a way to maximize the time spent in the atmosphere, so that the time of observation of the limb is larger than in the usual limb mode in the orbital plane [Bertaux et al., 2007a]. The signal recorded by the CCD detector is integrated over 640 ms periods typically during about 10 min. In this way, the line of sight scans a range of altitudes, generally between 70 km and 400 km and each second a full UV spectrum is obtained. For most of the observations reported here, the spacecraft was oriented in such a way that the SPICAV line of sight scanned the dark limb several times during the ascending portion of the VEX orbit. Since the line of sight crossed the same altitudes once during the ingress and once during the egress segments owing to the geometry of the grazing limb observation type, SPICAV supplied several sets of two (one for ingress, one for egress) altitude scans of five (for the 5 spatial bins) altitude profiles at each orbit. Consequently, as many as 6 5 limb profiles of the NO airglow were obtained on a same orbit between 6.1° and 71.8°N. At this date, a total of 17 orbits and 201 limb scans with suitable night airglow observations have been collected. As an example, Figure 1 shows the altitude variation of the altitude of the tangent point of the central point of the field of view for SPICAV spatial bins 1 and 5 during orbit 322. Other spatial bins scan the region indicated by the dotted gray area. Once all profiles are retrieved, the finite aperture of the field of view projected onto the limb is taken into account. For a given airglow layer, the apparent altitude of the emission peak and
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the maximum brightness depend on the value of the field of view projected on the limb. The angular field of view of one pixel is equal to 0.7 arcmin and the bin parameter varies between 2 and 32. The distance between the spacecraft and the observed tangent point lies within the 3000 to 12000 km range and the angle between the SPICAV slit and the limb is between 15° and 72°. For each observation, this set of parameters is used to determine the field of view projected on the limb. It varies from 3 to 27 km, but generally remains between 12 and 22 km, with a mean value of 14.9 km. We have investigated the effects of the smoothing by the field of view of an emission layer having a vertical Chapman profile integrated along the line of sight. Each Chapman profile is constrained to show the same topside scale height as the observation. It is found that the peak intensity observed by SPICAV decreases by a factor between 1 and 2 and the altitude of the peak is slightly above the unconvolved value. Data points in Figures 6 to 10 have been corrected for this smoothing effect by setting the peak intensity and altitude to the values they would have if the limb profiles had been observed with a negligibly small field of view. [10] For this study, we have integrated each spectrum over the entire d and g bands emission between 190 nm and 260 nm, including both the d and the g bands. As discussed in section 1, the total emission rate of the two molecular systems is equal to the N + O recombination rate. Figure 2 shows a sum of all (771) individual nightglow spectra obtained between 90 and 120 km on orbit 516. Comparison with the laboratory spectrum of the N(4S) + O(3P) recombination [Groth et al., 1971] confirms that this region of the Venus nightglow ultraviolet spectrum contains no other emission than those arising from the C 2P and A 2S states of NO. Integration of each individual spectrum over this spectral domain provides the total NO emission brightness in kR for each scanned altitude. Figure 3 shows limb profiles of the NO nightglow measured on ingress and egress segments of randomly selected orbit 271. In this plot, since the limb profiles in the five spatial bins are very similar as it is generally the case, the data points for spatial elements 1 and 2 have been grouped together. The data points were first binned in altitude (in 2 km steps), and these binned profiles were then co-added, separating egress and ingress data. Depending on altitude, each bin contains between 10 and 20 spectra, so that the error bar are quite small. The latitudes of the peak emission at ingress and egress are 21.7°N and 29.7°N and the local times are 2348 and 2330 LT, respectively. The maximum intensities at the peak are 46 and 35 kR, indicating the presence of a substantial gradient over a distance of only 8.0 degrees of latitude. The limb intensities have been numerically inverted using an Abel algorithm assuming a horizontally homogeneous distribution to calculate the corresponding vertical emission rates of 0.8 and 0.9 kR. These values may be compared with the 1.9 kR statistically observed by PVOUVS brightness in the bright spot. As just mentioned the limb profiles observed in the spatial bins are generally close to each other. However, strong horizontal gradients are occasionally present and noticeable differences between the profiles in the spatial bins may then be observed. An example is given in Figure 4 which shows four ingress smoothed profiles obtained during VEX orbit 322. The line of sight for spatial bin 5 did not intersect the altitude of the
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Figure 2. Spectrum of the nitric oxide ultraviolet nightglow obtained by summing all 771 individual spectra collected with the SPICAV instrument between 90 and 120 km on VEX orbit 516.
maximum of the airglow layer and is thus not represented in Figure 4. Moving from spatial bin 1 to 4, the limb intensity first increases then drops between adjacent bins angularly separated by only 5.6 arcmin. This separation maps into 1.8 (between bin 1 and 2), 1.4 (between bin 2 and 3) and 3.5 (between bin 3 and 4) degrees of increasing planetary latitude at the emission peak, which corresponds to 197, 158, and 374 km, respectively, of horizontal distance of the tangent points along the line of sight. Similarly, the altitude of the airglow peak varies from a bin to the next one. The differences in shape, peak altitude and brightness observed in the 4 spatial bins indicate that, in this case, the NO airglow distribution is significantly non uniform even over a restricted range of latitudes. It is important to note, however, that latitudinal gradients are usually much smaller, as was shown for example in Figure 3. [11] Each observed limb profile can be characterized by the altitude, brightness, latitude and local time of the maximum emission brightness. We first apply a filter to each limb profile to remove some of the high spatial frequency fluctuations through the data point following noise and background subtraction. From these fits we determine the peak altitude, the topside scale height of the emission as well as the total NO peak brightness. In this data set, covering the period between 17 January and 19 September 2007, a total of 201 good quality limb profiles were obtained. Their local time covers the range from 2000 to 0300 LT. The local time and latitudinal distribution of the observed limb profiles included in the database is represented in Figure 5. It shows that most observations were collected in two separate groups. Group 1 includes profiles between 6.1°N and 71.8°N observed in the post-midnight sector (2400 to 0300 LT), while group 2 which is more widespread corresponds to midlatitude to low-latitude observations between 2000 and 2200 LT. It should be noted that no observations were made in the
region of the NO bright spot observed with PV-OUVS south of the equator in the post-midnight period.
3. Correlations [12] We first examine the statistical distribution of the altitude of the peak of the airglow emission, corrected for the smoothing effect of the finite field of view. Figure 6 shows the distribution, where the data points have been distributed into 2.5 km altitude bins. The distribution is
Figure 3. Limb profile of the NO airglow measured on the ingress and egress of VEX orbit 271 on 17 January 2007. The latitudes of the peak emission are 21.7°N and 29.7°N and local times 2348 LT and 2330 LT for ingress and egress, respectively. The raw data are represented by open triangles and diamonds, while the smooth profile is obtained following Fourier filtering of high spatial frequencies. The error bars include the effect of the photon statistics and the detector thermal noise.
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Figure 4. Smoothed limb profiles observed during ingress of the line of sight during VEX orbit 327 in four adjacent spatial bins on the CCD detector. The error bars include the effect of the photon statistics and the detector thermal noise.
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Figure 6. Distribution of the altitudes of the tangent points corresponding to the peak in the airglow limb profiles analyzed in this study. The altitudes have been corrected to account for the smoothing effect of the SPICAV field of view (see text).
quite symmetric about the mean value, with some values above 125 km and a few profiles with a peak below 100 km. The mean value is 112.6 km, quite close to the median value of 112.9 km, with a 1-s deviation of 5.8 km. It should be noted that we implicitly assume a horizontal homogeneous layer. If the emission is not present (or weaker) at the limb but present (or stronger) in the foreground or in the background behind the limb, then the derived altitude of the peak is necessarily lower than the actual altitude. Therefore, this effect cannot explain the high-altitude peaks which are shown in Figure 6. [13] The distribution of the peak brightness with a 6 kR bin size is shown in Figure 7. It is much less symmetric, showing an extended tail of high-intensity values. The mean limb total NO brightness is 32 kR and the median value is 19 kR, with a standard deviation of 60 kR. The large emission rate of 440 kR observed at the limb on VEX orbit 516 (19 September 2007) has not been included in this nor in the following plots for reasons of graphical representa-
tions. The large scatter suggests that the large-scale winds which carry the O and N atoms from the dayside are intrinsically variable. As the wind pattern and velocity change, the amount of atoms transported to a given location and the position of the subsidence region considerably fluctuate, generating variability of the bright airglow spot location and intensity. [14] We now search for possible correlations between the altitude of the airglow, its brightness and latitude. Figure 8 shows the altitude of the airglow peaks plotted versus its latitude between 11° and 85°N. A slow decrease is observed toward the polar region as indicated by the linear regression obtained by minimizing the chi-squares, which shows a drop of about 5 km from low to high latitude. However, the linear correlation coefficient is equal to only 0.26. Figure 9 shows the lack of correlation (r = 0.19) between the
Figure 5. Local time and latitudinal coverage of the limb profiles used in the statistical study of the NO airglow distribution.
Figure 7. Distribution of the brightness of the peaks in the airglow limb profiles analyzed in this study. The intensities have been corrected to account for the smoothing effect of the SPICAV field of view.
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lation carries atoms from the day to the night side, where the combination of the global circulation and local turbulent transport carries the atoms in the region where they recombine near 115 km. The one-dimensional model developed for this study brings into play diffusion equations as well as production and loss by chemical reactions. It is similar to that described by Cox et al. [2008] to model the NO nightglow on Mars. The continuity equation for a minor constituent i in the thermosphere may be written @ni @f @ ðni wÞ ¼ i þ Pi L i @t @z @z
ð2Þ
with the vertical diffusive flux fi of a minor constituent i given by @ni ni @T Di K ni fi ¼ ðDi þ K Þ þ þ @z T @z Hi H
Figure 8. Relationship between the altitude of the tangent point of the emission peak and its latitude. airglow peak intensity and the latitude of the tangent point. However, the largest intensities are observed at latitudes between 20° and 50°N and 65° and 70°N. It should be noted that the large values between 65° and 70° are located far from the statistical bright spot observed with PV-OUVS. Figure 10 illustrates how these brightness data points are distributed with the altitude of the airglow peak. Although a considerable scatter is observed (r = 016), it appears that the large intensity limb profiles are generally concurrent with a peak altitude located in the vicinity of 113 ± 4 km, close to the mean altitude of the airglow layer. By contrast, profiles showing weaker emission rate correspond to situations where the peak altitude is outside this range of central values.
4. One-Dimensional Modeling [15] The use of a one-dimensional model makes it possible to evaluate the downward flux of nitrogen and oxygen atoms that recombine to produce the NO airglow emission. As described in the introduction, the current global view is that the subsolar to antisolar (SSAS) circu-
Figure 9. Relationship between the brightness of the emission peak and its latitude.
ð3Þ
with Di the molecular diffusion coefficient for constituent i, K is the vertical eddy diffusion coefficient, Hi is the local scale height of the ith constituent. H is the atmospheric scale height, T is the neutral gas temperature, ni is the number density of the ith species, z is the altitude, t is time, Pi is the production rate of species i, Li the loss rate and w is the vertical velocity positive upward. The last term in equation (2) corresponds to the vertical advective flux. We adopt the vertical variation of the eddy diffusion coefficient K similar to that used in earlier studies [von Zahn et al., 1979; Ge´rard et al., 1981], that is, A K ð zÞ ¼ pffiffiffiffiffiffiffiffiffi cm2 s1 nð zÞ
ð4Þ
where n is the total number density and A is a free parameter of the model which is independent of altitude. Therefore, K only depends on altitude and is identical for all constituents. The model solves equation (2) for O(3P), N(4S), NO and O2 (1D). The Pi and Li coefficients depend on the choice of the different reactions that come into play and which are listed in Table 1 with their corresponding reaction coefficients. The temperature dependence of the
Figure 10. Relationship between the tangent point altitude of the emission peak and its brightness.
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Table 1. Chemical Reactions and Rate Coefficients Rate Coefficienta (cm3 s1)
Reaction ðR1Þ ðR2Þ ðR3Þ ðR4Þ
N þ O ! NO þ hv N þ O þ CO2 ! NO þ CO2 N þ NO ! N2 þ O O þ O þ CO2 ! O2 þ CO2
1.9 2.8 2.5 2.8
1017 (300/T)1/2 (1 – 0.57/T1/2) 1032 (300/T)1/2 1010 (T/300)1/2 exp(600/T) 1032
Reference Dalgarno et al. [1992] Campbell and Thrush [1966]b Fox [1994] Campbell and Gray [1973]c
Except for three-body recombination reactions (R2) and (R4) in cm6 s1. The rate coefficient of (R2) is the value measured with M = N2 as a third body, multiplied by 2.5 for M = CO2 with an inverse square root dependence on temperature. c The rate coefficient is the value measured for M = N2 at 196 K multiplied by 2.5, a factor recommended as appropriate for Mars atmosphere models [Nair et al., 1994]. a
b
N + O + CO2 reaction was introduced by Stewart et al. [1980] and used in subsequent models such as Yung and DeMore [1982]. For reasons of continuity, we have adopted the same temperature dependence as earlier studies. Equation (2) is solved numerically between 130 and 80 km using the finite volume method on a constant grid. For O and N, we apply a flux boundary condition through the upper boundary, and assume photochemical equilibrium at the lower boundary and we use density null vectors for initial conditions (except at the boundaries). The model is then parameterized by A, w(z), Fo and FN. In fact, the choice of the densities at 80 km is relatively arbitrary and it does not influence the solution further than one scale height from the lower boundary if it is kept in a reasonable range. At the upper boundary, we leave the O(3P) and the N(4S) fluxes as free parameters to be determined by fitting the modeled limb profile to the observations. The limb profile of the nitric oxide emission is obtained by integrating the k[O][N] product along the line of sight, where k is the rate coefficient of the O + N recombination. We note that in all limb profiles most of the emission is located below 135 km, the altitude of the homopause determined from the composition measurements made by the Pioneer Venus large probe during its descent [von Zahn et al., 1979]. Onedimensional models usually do not include advection terms. Consequently, in this representation, vertical transport is solely the result of the molecular and eddy diffusions, and
advection is implicitly contained in the K coefficient, despite the lack of physical meaning. [16] We first present results of modeled limb profiles using CO2 density and temperature taken from the Venus Reference Atmosphere (VIRA) model in the code formulation given by Hedin et al. [1983]. Figure 11 shows a comparison between three observed limb profiles and those simulated using the one-dimensional model described above, following integration of the calculated volume emission rate along the line of sight. Absorption of the NO bands by CO2 is less than 1% in the altitude region considered here and is thus negligible. The simulated distributions match reasonably well the peaks of the observed limb profiles. As in earlier studies [Stewart et al., 1980; Ge´rard et al., 1981], we have assumed that the flux of N(4S) is 1% of the O flux, in agreement with the average O/N number density ratio measured in the dayside thermosphere with the neutral mass spectrometer on board Pioneer Venus. The observational data have been smoothed as described in section 2 to remove most of the scatter associated with instrument noise. In Figure 11, three comparisons are made with profiles that have been selected to cover a range of altitudes of the emission peak. The profile for orbit 320 obtained on 7 March 2007 at 32.3°N and 0036 LT shows a maximum at 112 km and is typical of a vertical distribution frequently observed. The peak intensity reaches 39 kR, a value larger than the median but close to
Figure 11. Examples of airglow limb profiles with different peak altitudes (circles). The solid lines show the corresponding model simulations that best fit the altitude and brightness of the airglow maximum. 7 of 10
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Figure 12. Distribution of the apparent topside scale heights of the airglow limb profiles in the region located above the emission peak. For comparison, the dotted line indicates the airglow scale height for diffusive equilibrium of O and N, while the dashed line is the expected value for a downward flow of atoms in the homosphere (see text). the mean value, as was shown in Figure 7. The calculated limb profile represented by the solid line was obtained using the CO2 density and temperatures profiles from the Venus Reference Atmosphere (VIRA) model in the code formulation given by Hedin et al. [1983], for the appropriate latitude, local time and solar activity level. The calculated limb profile has been convolved by the value of the field of view projection on the planet calculated for the pixel binning and the spacecraft-Venus distance at the time of the observation. The parameters used for this simulation were FO = 2.1 1011 cm2 s1, FN = 2.1 109 cm2 s1, A = 7.4 1011. The altitude distribution at the peak and the topside scale height are well reproduced by the model, whereas the calculated bottomside scale heights are somewhat larger than the observed values. Bertaux et al. [2007b] showed that SPICAV stellar occultations by CO2 in the Venus thermosphere suggest that temperatures are as much as 40 K above the VIRA values in the region of the antisolar point. The effect of this different temperature distribution has been assessed by substituting the CO2 and temperature profiles observed by Bertaux et al. on orbit 104 (near midnight) to the VIRA distribution. The shape and absolute values of the predicted intensity remains almost unaffected when the warmer distributions deduced from the SPICAV occultation are used. [17] To document the question of the scale height, for each limb profile, we have determined the value of the apparent emission scale height H in the region extending between Zp + 5 km and Zp +15 km (except if the scale height varies rapidly in this altitude range), where Zp is the altitude of the emission peak. The resulting distribution is shown in Figure 12. The distribution of H values covers more than an order of magnitude, ranging from about 2 to more than 30 km. If the homopause is located above the airglow peak region, as suggested by the determination of the homopause altitude from the Pioneer large probe descent, the airglow layer is in the homosphere dominated
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by eddy mixing. In these conditions, the analytical solution to equations (2) and (3) indicates that the topside scale height of the constituents flowing downward toward the chemical sink region varies as 1/K. With the altitude dependence adopted in our model, this implies that both the O and N densities vary as the square root of the CO2 density. Consequently, since the airglow volume emission rate is proportional to the product of the O and N densities, the airglow scale height above the peak altitude is close to the value of the background atmosphere (mostly CO2), which, according to the VIRA model, is on the order of 3.1 km at 0000 LT near the equator. Instead, if the emission is located above the homopause and controlled by molecular diffusion in the absence of wind, the airglow scale height is that of the geometric mean 4.6 km of the O and N individual scale heights. Both values are indicated as vertical bars in Figure 12. The scale height determined from the analysis of the limb scans observed closer to the planet with PV-OUVS by Ge´rard et al. [1981] was 3.2 km, in close agreement with the value expected for turbulent downward transport of the atoms. It is clear that a large fraction of the observations (over 90%) exhibit an apparent scale height in excess of this value. Most of these larger values are observational effects due to the combination of (1) a finite field of view of the instrument whose value depends on the spatial bin size used for a given observation, (2) the distance between the spacecraft and (3) the slit orientation relative to the planetary limb. As mentioned before, the projected aperture of the field of view on the limb ranges between 3 and 27 km, which tends to smooth the observed limb profile and to increase the apparent topside scale height. [18] Figure 13 shows the O(3P) and N(4S) vertical distribution calculated from the best fit to the limb profile of orbit 320 in Figure 11. The maximum atomic oxygen density is predicted at 111 km where it reaches 4 1011 cm3. This value is in good agreement with the O density determined by Drossart et al. [2007] and Ge´rard et al. [2008] from the vertical distribution at the limb of the O2(1D) infrared airglow observed with the VIRTIS spectral imager. The atomic nitrogen density profile presents a peak of 2 108 cm3 at 122 km. This result may be compared with the N density profiles calculated previously by Stewart et al. [1980] or by Bougher et al. [1990] for solar maximum activity at 2400 LT with the Venus TGCM and VIRA models and reported in their Figure 5. The calculated maximum N density was about 7 108 cm3 at 112 km, in reasonably good agreement with the result of Figure 13, considering the different solar activity levels and local times. In Figure 11, two other simulations are compared with observed NO limb profiles. One (orbit 321, lat 53.8°N, LT = 0048) shows a peak at 105 km and is very well reproduced by the 1-D model both in magnitude and shape. The last one (orbit 324, lat 17.9°N, LT = 0118) presents a shallow peak at 119 km, while the best fit model produces maximum limb intensity at 112 km. The parameters of the fit for orbit 321 are fO = 6.6 1010 cm2 s1, A = 4 1012. For orbit 324, the parameters are fO = 8.2 1010 cm2 s1, and A = 0. In this latter case, downward transport solely by molecular diffusion alone is even slightly too strong to match the emission peak at 119 km. This result suggests that eddy diffusion may be locally quite weak or that upward
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Figure 13. Vertical distribution of the atomic oxygen and nitrogen number density obtained from the model fitting the limb profile of orbit 320 shown in Figure 11. vertical winds may act to raise the altitude of the airglow layer. The observed airglow topside scales heights are 13.3 km (orbit 320), 16 km (321) and 14.5 km (324). All three simulated profiles smoothed by the appropriate projected field of view exhibit an apparent scale height in agreement with the observed intensity drop above the emission peak. The modeled emission scale heights before smoothing are 3.6 km, 3.7 km and 4.6 km for orbits 320, 321 and 324, respectively. These values lie between the diffusive and the turbulent scale heights of 4.6 km and 3.1 km as discussed before.
5. Conclusions [19] The NO nightglow emission is a good tracer of atmosphere dynamics in the Venus thermosphere. The nightside lower thermosphere of the planet remains relatively unexplored, especially the minor constituents densities and their dynamics. SPICAV observations of the NO nightglow provide a set of limb profiles data that have been used to analyze the NO vertical distribution in detail. From the comparison with a one-dimensional chemicaldiffusive model, we have determined values for the eddy diffusion coefficient. The value of the A coefficient in formula (4) is 7.4 1011, 4 1012 and 0 for the cases modeled in this study. The first two values are 0.09 to 0.5 smaller than the earlier determination based on NO airglow observations from Pioneer Venus. We note that the PVOUVS limb observations were made at low latitudes near maximum solar activity. The profiles described here were obtained in the northern hemisphere, generally away from the bright spot reported by Stewart et al. [1980]. The nitrogen downward fluxes at 130 km range between 1 108 and 4 109 cm2 s1 with typical values of 2 109 cm2 s1. The deduced O fluxes are less than the average column production rate of O atoms on the Venus dayside which was estimated at 8 1012 cm2 s1 by Leu and Yung [1987], in agreement with the discussion by Ge´rard et al. [2008]. Similarly, the daytime N average column production was found to be 1.3 1010 atoms cm2 s1 by Ge´rard et al. [1988] at solar maximum and
about 9 109 atoms cm2 s1 in the VTGCM photochemistry [Bougher et al., 1990]. The N vertical fluxes on the nightside derived from these observations thus also appear compatible with the availability of N atoms on the dayside. We also find that the topside scale height of the NO airglow layer is in agreement with that expected for constituents in diffusive equilibrium or diffusing downward through the homosphere. [20] One of the main results in the Venus Express observations is the large variability in the brightness of the NO nightglow emission. This feature was already apparent in the individual NO brightness maps obtained from PV-OUVS presented by Stewart et al. [1980] and in Figure 9 by Bougher et al. [1990]. It is not related to concurrent solar activity which remained very low throughout the period of these observations. It implies timescales as short as an Earth day or less. A similar variability is present in the O2 (1D) emission arising from three-body recombination of O atoms at 1.27 mm that have been observed with VIRTIS-M, also on board Venus Express. Both limb and nadir observations [Drossart et al., 2007; Ge´rard et al., 2008; Hueso et al., 2008; G. Piccioni et al., Molecular oxygen nightglow from the VIRTIS near IR observations in the Venus upper mesosphere, submitted to Journal of Geophysical Research, 2008] indicate that the brightness of the O2 emission and its peak altitude exhibit substantial variations sometimes on timescales on the order of an hour or so. This variability may be caused by time variations in gravity waves breaking which potentially decelerate the SSAS flow. [21] The SPICAV limb observations confirm and complete this picture and they also show the presence of strong horizontal gradients. Such gradients may be interpreted as signatures of localized regions of enhanced or decreased downward fluxes of atoms producing the patchy structure that had been observed on individual maps. In the southern hemisphere, the larger distance of the spacecraft from the planet precludes high-resolution limb observations. Therefore, the statistical bright spot near midnight slightly south of the equator was generally not captured in this tangential limb observing mode. It is thus not possible at this stage to
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confirm whether the emission maximum is at the same location during solar minimum periods. If so, the comparison with the morphology of the O2 (1D) emission due to three-body recombination of O atoms which has a peak near midnight near the equator would suggest that the zonal winds which shift the NO airglow at 113 km toward dawn have a different regime near 96 km. In any case, the features described in this study should be additional constraints to future three-dimensional models of the Venus nightside. [22] Acknowledgments. The authors thank the SPICAV and the Venus Express teams for the excellent quality of their work. J.-C. Ge´rard is supported by the Belgian Fund for Scientific Research (FNRS). This research was funded by the PRODEX program of the European Space Agency (ESA) managed with the help of the Belgian Space Policy Office. The construction of the SPICAV instrument was funded by CNRS, CNES, and ESA/PRODEX.
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Feldman, P. D., H. W. Moos, J. T. Clarke, and A. L. Lane (1979), Identification of the UV nightglow from Venus, Nature, 279, 221 – 222, doi:10.1038/279221a0. Fox, J. L. (1994), Rate coefficient for the reaction N + NO, J. Geophys. Res., 99, 6273 – 6276, doi:10.1029/93JA03299. Ge´rard, J.-C., A. I. F. Stewart, and S. W. Bougher (1981), The altitude distribution of the Venus ultraviolet airglow and implications on vertical transport, Geophys. Res. Lett., 8, 633 – 636, doi:10.1029/GL008i006p00633. Ge´rard, J.-C., E. J. Deneye, and M. Lerho (1988), Sources and distribution of odd nitrogen in the Venus daytime thermosphere, Icarus, 75, 171 – 184, doi:10.1016/0019-1035(88)90135-2. Ge´rard, J.-C., A. Saglam, G. Piccioni, P. Drossart, C. Cox, S. Erard, R. Hueso, and A. Sa´nchez-Lavega (2008), The distribution of the O2 infrared nightglow observed with VIRTIS on board Venus Express, Geophys. Res. Lett., 35, L02207, doi:10.1029/2007GL032021. Groth, W., D. Kley, and U. Schurath (1971), Rate constant for the infrared emission of the NO (C2P ! A2S+) transition, J. Quant. Spectrosc. Radiat. Transfer, 11, 1475 – 1480, doi:10.1016/0022-4073(71)90109-9. Hedin, A. E., H. B. Niemann, W. T. Kasprzak, and A. Seiff (1983), Global empirical model of the Venus thermosphere, J. Geophys. Res., 88, 73 – 83. Hueso, R., A. Sa´nchez-Lavega, G. Piccioni, P. Drossart, J. C. Ge´rard, I. Khatuntsev, L. Zasova, and A. Migliorini (2008), Morphology and dynamics of Venus oxygen airglow, J. Geophys. Res., 113, E00B02, doi:10.1029/2008JE003081. Leu, M.-T., and Y. L. Yung (1987), Determination of O2 (a1Dg) and O2 (b1S+g) yields in the reaction O + ClO ! Cl + O2: Implications for photochemistry in the atmosphere of Venus, Geophys. Res. Lett., 14(9), 949 – 952, doi:10.1029/GL014i009p00949. Nair, H., M. Allen, A. D. Anbar, Y. L. Yung, and R. T. Clancy (1994), A photochemical model of the Martian atmosphere, Icarus, 111, 124 – 150, doi:10.1006/icar.1994.1137. Stewart, A. I., and C. A. Barth (1979), Ultraviolet night airglow of Venus, Science, 205, 59 – 62, doi:10.1126/science.205.4401.59. Stewart, A. I. F., J.-C. Ge´rard, D. W. Rusch, and S. W. Bougher (1980), Morphology of the Venus ultraviolet night airglow, J. Geophys. Res., 85, 7861 – 7870, doi:10.1029/JA085iA13p07861. Svedhem, H., et al. (2007), Venus Express-The first European mission to Venus, Planet. Space Sci., 55, 1636 – 1652, doi:10.1016/j.pss.2007.01.013. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017. von Zahn, U., K. H. Fricke, H. J. Hoffmann, and K. Pelka (1979), Venus: Eddy coefficients in the thermosphere and in the inferred helium content of the lower atmosphere, Geophys. Res. Lett., 6, 337 – 340, doi:10.1029/ GL006i005p00337. Yung, Y. L., and W. B. DeMore (1982), Photochemistry of the stratosphere of Venus: Implications for atmospheric evolution, Icarus, 51, 199 – 247, doi:10.1016/0019-1035(82)90080-X.
J.-L. Bertaux, C. Nehme´, and E. Villard, Universite´ de Versailles SaintQuentin-UVSQ, Service d’Ae´ronomie du CNRS, Boite Post 3, F-91371 Verrie`res-le-Buisson, France. C. Cox, J.-C. Ge´rard, and A. Saglam, Laboratoire de Physique Atmosphe´rique et Plane´taire, Universite´ de Lie`ge, Baˆtiment B5c, Alle´e de 6 Aouˆt, 17, B-4000 Lie`ge, Belgium. (
[email protected])
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Concurrent observations of the ultraviolet nitric oxide and infrared O2 nightglow emissions with Venus Express J.-C. Ge´rard,1 C. Cox,1 L. Soret,1 A. Saglam,1 G. Piccioni,2 J.-L. Bertaux,3,4 and P. Drossart5 Received 19 February 2009; revised 4 May 2009; accepted 18 June 2009; published 22 September 2009.
[1] Two prominent features of the Venus nightside airglow are the nitric oxide d and g
bands produced by radiative association of O and N atoms in the lower thermosphere and the O2 infrared emission generated by three-body recombination of oxygen atoms in the upper mesosphere. The O2 airglow has been observed from the ground, during the Cassini flyby, and with VIRTIS on board Venus Express. It now appears that the global structure of the two emissions shows some similarities, but the statistical location of the region of strongest emission is not coincident. The Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV) ultraviolet spectrograph has collected a large number of spectra of the Venus nitric oxide nightside airglow. Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) images have been obtained at the limb and in the nadir-viewing mode and have provided new information on the horizontal and vertical distribution of the emission. We present the first concurrent observations of the two emissions observed with Venus Express. We show that nadir observations generally indicate a low degree of correlation between the two emissions observed quasi-simultaneously at a common location. A statistical study of limb profiles indicates that the altitude and the brightness of the two airglow layers generally do not covary. We suggest that this lack of correlation is explained by the presence of strong horizontal winds in the mesosphere-thermosphere transition region. They carry the downflowing atoms over large distances in such a way that regions of enhanced NO emission generally do not coincide with zones of bright O2 airglow. Citation: Ge´rard, J.-C., C. Cox, L. Soret, A. Saglam, G. Piccioni, J.-L. Bertaux, and P. Drossart (2009), Concurrent observations of the ultraviolet nitric oxide and infrared O2 nightglow emissions with Venus Express, J. Geophys. Res., 114, E00B44, doi:10.1029/2009JE003371.
1. Introduction [2] Planetary airglow is a powerful way to remotely probe the characteristics of atmospheres from an orbiting or flying-by spacecraft that complements in situ measurements. In particular, the study of airglow morphology, time variations and brightness may provide key observations concerning atmospheric composition, temperature structure, transport processes and their response to the solar photon and particle inputs. In the case of Venus, the Venera and Pioneer Venus missions have shed new light on some of these aspects more than 25 years ago. The presence of the delta and gamma bands of nitric oxide in the Venus nightglow was detected and identified by Feldman et
al. [1979] using the ultraviolet spectrograph on board the International Ultraviolet Explorer (IUE). It was also observed by Stewart and Barth [1979] with the ultraviolet spectrometer on board the Pioneer Venus Orbiter (PVOUVS). The emission is produced by radiative recombination through inverse predissociation of nitrogen N(4S) and oxygen O(3P) atoms and dominates the middle ultraviolet nightglow spectrum. In this process, excited NO molecules radiate the ultraviolet d and g bands between 180 and 310 nm: N þ O ! NO C 2 P
giving rise to NO C 2 P ! NO X 2 P þ d-bands
1
Laboratoire de Physique Atmosphe´rique et Plane´taire, Universite´ de Lie`ge, Lie`ge, Belgium. 2 INAF, IASF, Rome, Italy. 3 LATMOS, Verrie`res-le-Buisson, France. 4 Institut Pierre Simon Laplace, Universite´ de Versailles Saint Quentin, Guyancourt, France. 5 LESIA, Observatoire de Paris, Meudon, France. Copyright 2009 by the American Geophysical Union. 0148-0227/09/2009JE003371$09.00
NO A 2 S; v0 ¼ 0 ! NO X 2 P þ g-bands
Emission from the C 2P (v = 0) ! A 2S (v = 0) transition at 1.224 mm was recently observed with the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS) -M in the
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Venus nightglow [Garcı´a Mun˜oz et al., 2009]. The N and O atoms are mainly produced by dissociation of N2 and CO2 on the dayside by extreme ultraviolet (EUV) photons and photoelectrons. [3] Stewart et al. [1980] obtained images of the Venus nightside in the d (0,1) band at 198 nm every 24 h with PVOUVS when Pioneer Venus was near apoapsis. They showed that the emission was highly variable in brightness and morphology over consecutive 24 h periods. The location of the brightest spots ranged from 2130 to 0300 LT and 39°S to 60°N [Bougher et al., 1990]. This variability appeared to be caused by instabilities in the large-scale circulation, possibly as a result of wind shears near the terminator or time-varying wave drag from gravity waves. Stewart et al. [1980] built up a statistical global map of the UV nightglow showing that the emission is concentrated in a bright spot located near 0200 local solar time, south of the equator. Bougher et al. [1990] estimated the emission rate of this enhanced emission to be 1.9 kilorayleighs (kR) whereas the average hemispheric nightside intensity is 0.48 kR. These observations confirmed the general picture where production of O and N atoms by solar EUV on the dayside is followed by global circulation to the nightside, downward transport and radiative recombination. The shift of the statistical bright spot toward dawn was interpreted as a signature of a residual superrotation into the lower thermosphere. A determination of the altitude of the emission peak by Ge´rard et al. [1981] concluded that the emission peak is located close to 115 km. Using a one dimensional model, they derived an eddy diffusion coefficient K 8 1012 n1/2 cm2 s1, where n is the total number density. The general picture of production of O and N atoms followed by transport to the nightside by the subsolar to antisolar circulation, downward turbulent mixing and radiative recombination appeared quantitatively consistent with the PV-OUVS observations. This concept was numerically validated by three-dimensional simulations using the Venus Thermospheric General Circulation Model (VTGCM) [Bougher et al., 1990]. The statistical location of the bright spot was reasonably well predicted by the threedimensional model and implied zonal winds of about 50– 75 m s1 in the 115 –150 km region. The observed shift toward dawn of the statistical location of the airglow maximum was reproduced by the VTGCM. [4] Limb observations of the spectrum of nightglow emission in the d and g bands of NO with the Spectroscopy for Investigation of Characteristics of the Atmosphere of Venus (SPICAV) ultraviolet spectrometer on board Venus Express [Titov et al., 2006; Svedhem et al., 2007] have been recently reported by Ge´rard et al. [2008a]. The mean altitude of the emission layer was found to be located at 113 km, with variations between 95 and 132 km. The mean limb brightness of the total NO emission at the limb was 32 kR, but it is highly variable with limb intensities as large as 440 kR and values below 5 kR at northern midlatitudes. It was found that the mean altitude of the emission peak statistically drops with increasing latitude between 6° and 72°N. From model fits to observed profiles, they determined that the downward flux of N atoms at 130 km typically varies between 1 108 to 4 109 atoms cm2 s1. The eddy diffusion coefficient K deduced from comparisons to the observed limb profiles was significantly less than that
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determined from the observations made with the Pioneer Venus UV spectrometer at low latitudes during periods of high solar activity. [5] The oxygen airglow (0– 0) emission at 1.27 mm is the most intense nonthermal component in the Venus atmosphere. It belongs to the a1Dg – X 3Sg Atmospheric Infrared system and corresponds to an electric dipole forbidden transition with a radiative lifetime of about 70 min [Miller et al., 2001]. It was first discovered in ground-based observations of Venus by Connes et al. [1979] and subsequently imaged with ground-based telescopes [Alien et al., 1992; Crisp et al., 1996; Lellouch et al., 1997; Ohtsuki et al., 2008; Bailey et al., 2008]. It was measured from space during the Cassini flyby with a local maximum brightness of about 4 MR [Drossart et al., 1993]. The oxygen IR nightglow appeared patchy, highly variable with asymmetries, often exhibiting multiple local maxima, with variations on time scales as short as 1 h. A much weaker O2 airglow was also detected in the Herzberg II visible wavelengths [Krasnopolsky, 1986; Bougher and Borucki, 1994]. The 1.27 mm emission is produced by recombination of oxygen atoms created by photodissociation of CO2 and CO at thermospheric altitudes on the dayside. As previously explained, O atoms are transported to the nightside by the global thermospheric circulation. Three-body recombination of O atoms in the upper nightside mesosphere (95 –110 km) leads to O2 formation in excited states, followed by airglow emissions as the molecules relax to the X 3Sg ground state: O þ O þ M ! O2 * þ M
followed by O2 * ! O2 þ hn
where O2* indicates one of the excited states of the O2 molecule and M is any neutral constituent. A fraction of the O2 molecules, estimated to be about 7%, is formed directly in the a1Dg metastable state. A substantial fraction of the upper excited states cascades into the 1D state, so that the net efficiency of the production of this state in the threebody recombination may be close to 100% [Crisp et al., 1996]. Below the emission peak, O2 1Dg molecules may be deactivated by collisions with CO2, causing nonradiative transitions to the O2 ground state. The altitude of the peak of the 1.27 mm emission is thus controlled by the competition between vertical transport, recombination and quenching of O atoms. The emission rate is related to the downward flux of oxygen atoms. [6] The VIRTIS-M O2 airglow limb observations have been presented by Drossart et al. [2007a], Ge´rard et al. [2008b, 2009], and Piccioni et al. [2009a]. Drossart et al. [2007a] determined that the O2 peak emission is located near 96 km, which is consistent with three-body recombination of oxygen atoms. Ge´rard et al. [2008b] found that limb profiles observed at northern midlatitudes exhibit large intensity variations over short time periods. Ge´rard et al. [2009] described further emission limb profiles extracted from the images. They determined the vertical distribution of O2 (1Dg) atoms using an Abel inversion of the radiance limb profiles. Assuming photochemical equilibrium for O2 (1D), they used these density profiles combined with the CO2
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vertical distribution to determine the atomic oxygen density. Piccioni et al. [2009a] analyzed limb measurements from 42 orbits. They found that the peak altitude of the O2 (1Dg) volume emission rate is typically located between 95 and 100 km, with a mean value of 97.4 ± 2.5 km. The vertical profile is broader near the equator, with a full width at half maximum of 11 km, a factor 2 larger than at middle latitudes. They reported that a secondary peak is frequently observed between 103 and 105 km. [7] VIRTIS nightside observations from Venus Express have complemented ground-based observations at much higher spatial resolution. In addition, limb observations from an orbit around Venus have given unprecedented access to the vertical distribution of the airglow layer and provided key constraints on the models. Drossart et al. [2007a] confirmed that the O2 nightglow exhibits a large spatial and temporal variability. Observations by VIRTIS in the nadir mode have been used to construct extensive maps of the Venus atmosphere in the O2 emission band [Ge´rard et al., 2008b; Piccioni et al., 2009a]. In nadir viewing geometry, the contamination of the O2 airglow by the thermal emission of the deeper atmosphere has to be subtracted to obtain clean O2 airglow images. The mean value, integrated over the nightside of the southern hemisphere, is typically about 0.8 MR, which is in agreement with the early groundbased observations giving a mean brightness of 1.2 MR for the night side. Hueso et al. [2008] found that the airglow is highly inhomogeneous with the regions of highest intensity generally located at low latitudes near the midnight meridian. They showed that zonal velocity derived from the motion of airglow features is dominated by an intense prograde jet from dawn to midnight extending up to 22 h in local time, with lower velocities and reversed sign from dusk. The brightest small-scale (100 km) features appeared correlated with locations of apparent convergence which may be a signature of compression and downwelling. Piccioni et al. [2009a] described the characteristics of the horizontal distribution of the airglow and showed that regions of high O2 airglow intensity are associated with downwelling causing an increase of the infrared brightness temperature. A similar conclusion was reached by Bailey et al. [2008], who derived rotational temperatures in excess of the VIRA values in regions of enhanced O2 1D emission rate. They associated these regions with conditions of larger downflow velocities where local temperature is increased by compressional heating. Similar and larger temperature enhancements were observed with SPICAV from measurements of UV CO2 absorption measurements during stellar occultations by Bertaux et al. [2007a]. The nonhomogeneous, time-dependent distribution of the O2 1Dg nightglow indicates that the local downward flow of oxygen may differ substantially from the mean value, in response to variations in the efficiency of the global day-to-night transport, the focusing effect of the nightside subsidence, changing zonal wind speeds, eddy transport efficiency, and gravity wave breaking [Bougher et al., 2006]. [8] The NO and O2 nightglows do not occur at the same altitude and thus provide information about different vertical levels: 95 – 105 km for O2 airglow [Piccioni et al., 2009a] and 115 km for NO [Ge´rard et al., 2008a]. In this study, we address the question of the covariation of the two airglow layers. Earlier studies have established that the
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statistical location of the nightglow bright spots is not coincident. This result was unexpected since the two emissions are produced by recombination of atoms created on the dayside by photodissociation and transported to the nightside by the subsolar to antisolar global circulation. In this study, we take advantage of the unique opportunity offered by two instruments of the Venus Express mission (SPICAV and VIRTIS) to observe almost simultaneously the two emissions and to investigate if their characteristics covary in time and space. The observations reported here were obtained both in nadir-viewing geometry, where horizontal variations of the emission rate can be mapped, and at the limb, where vertical variations are best investigated. We describe both types of observations and draw conclusions on the level of covariance we have observed and transport processes occurring in the Venus mesospherethermosphere transition region.
2. Parallel Nadir Observations [9] The Venus Express spacecraft moves along a quasipolar eccentric orbit with a 24-h period. The apocenter is located at 66,000 km, while the altitude of the pericenter (at 80°N) has varied between 250 km and 185 km. The orbit is fixed in the inertial space and therefore precesses at the rate of 1.6° d – 1. The precession of the orbital plane leads to a wide variety of configurations on the nightside as well as on the dayside. Several observation modes (science cases) may be selected including nadir observations, star pointing for stellar occultations by Venus’ atmosphere, fixed point tracking and limb observations [Titov et al., 2006]. [10] The SPICAV instrument and its performances were described by Bertaux et al. [2007b]. The ultraviolet spectrometer covers a spectral range extending from 118 nm to 320 nm including the NO d (C 2P X 2P) and g (A 2S X 2P) emission bands, the only spectral features with Lymana observed in the Venus nightglow [Ge´rard et al., 2008a]. The detector is a 407 288 pixel CCD and the angular field of view of one pixel is equal to 0.7 0.7 arcmin. For reasons of telemetry limitations and because of the time needed to read all the lines of the CCD, only 5 adjacent zones of the CCD detector are usually read out. In these nadir observations, the width of each spatial bin is 32 pixel lines, corresponding to a field of view of 3.7°. These lines are seen through the large (500 mm) slit, providing a spectral resolution of about 12 nm. The planetary area intercepted by the field of view depends on the location on its orbits. The spacecraft altitude ranged between 7350 and 9050 km during the nadir observations reported here. The SPICAV CCD is read out every second, but the actual integration period of each spectrum is 640 ms. The nonuniform dark current and offset values are carefully subtracted in each individual spectrum, using similar observations performed with a null amplification. The absolute calibration obtained by observing well-known hot stars spectra is then applied to obtain nitric oxide emission rates in kR [Bertaux et al., 2007b]. [11] Spectral images have been regularly obtained in nadir geometry with VIRTIS mostly from segments of the orbit near apocenter. The VIRTIS [Drossart et al., 2007b; Piccioni et al., 2009b] pixel size of 0.25 mrad gives a spatial resolution of 15 km on Venus from apocenter. For this
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Table 1. Time, Location, Maximum Brightness and Correlation Coefficient R of the Nitric Oxide and O2 (1D) Airglow Concurrent Observations by SPICAV and VIRTIS-M at Nadir
Orbit
Start Time (UT)
Stop Time (UT)
Local Time
Minimum Latitude (deg)
243 341 342 343 345 346 453 459 567 571 592 901 905 907
20 Dec 2006; 0752:46 28 Mar 2007; 0549:01 29 Mar 2007; 0547:05 30 Mar 2007; 0545:17 1 Apr 2007; 0542:15 2 Apr 2007; 0540:55 18 Jul 2007; 0710:30 24 Jul 2007; 0650:06 9 Nov 2007; 0055:00 13 Nov 2007; 0057:30 4 Dec 2007; 0213:35 8 Oct 2008; 0811:21 12 Oct 2008; 0819:58 14 Oct 2008; 0824:18
20 Dec 2006; 0806:29 28 Mar 2007; 0558:43 29 Mar 2007; 0606:39 30 Mar 2007; 0557:22 1 Apr 2007; 0556:04 2 Apr 2007; 0605:27 18 Jul 2007; 0735:42 24 Jul 2007; 0718:12 9 Nov 2007; 0123:40 13 Nov 2007; 0121:04 4 Dec 2007; 0231:45 8 Oct 2008; 0830:11 12 Oct 2008; 0838:48 14 Oct 2008; 0843:08
0035 2307 2314 2318 2330 2336 2306 2343 2312 0010 0156 2301 2328 2339
1.8 9.3 10.2 10.9 11.7 12.0 16.4 18.8 10.9 16.4 13.4 13.1 12.8 12.7
study, we use the VIRTIS-M mode which provides spectral cubes between 0.25 and 5 mm at a spectral resolution R 200. Each spectral channel is 9.5 nm wide in the region of the O2 Atmospheric Infrared system emission. A spatial scan, covering a 64 mrad 64 mrad field of view is obtained using a scanning mirror. However, even from apocenter, only a fraction of the Venus disk is observed during a mirror scan of the instrument and a spacecraft repointing is needed to collect a more extended coverage. For each VIRTIS image, the thermal contribution from the lower atmosphere is subtracted from the total signal using the VIRTIS fluxes measured in the three adjacent channels centered on 1.27 mm. The count rate is expressed in radiative flux units and MR using the measured instrumental calibration and the O2 (1Dg) relative line intensity for a temperature of 200 K. Airglow radiation emitted downward and subsequently backscattered by the underlying clouds is
Maximum Latitude (deg)
Imax NO (kR)
Imax O2 (MR)
R
48.9 5.1 10.9 0.4 2.3 14.1 38.1 37.5 31.6 19.0 48.3 20.2 20.2 20.4
2.7 1.0 3.1 1.0 2.1 3.7 1.1 2.2 3.5 2.5 6.2 1.7 2.5 3.9
1.1 0,6 1.1 2.2 0.9 3.5 0.6 0.9 1.1 1.1 1.4 0.8 0.7 0.8
0.55 0.84 0.41 0.74 0.23 0.32 0.13 0.08 0.86 0.59 0.25 0.44 0.49 0.26
accounted for using the correction factor derived by Crisp et al. [1996]. [12] The SPICAV and VIRTIS databases have been searched to identify periods when the fields of view of the nadir-viewing observations of both instruments overlapped over a significant time span. Table 1 lists the orbit numbers, times, and locations of these occurrences. As an example, Figure 1 illustrates the spatial coverage in the Venus atmosphere of the SPICAV slit (in gray) and the VIRTIS images (in black) during orbit 243. As can be seen, a common region was observed by both instruments northward of 2°N and southward of 8°S. The shape of the VIRTIS image coverage is defined by the combination of the decreasing spacecraft altitude during the 688 s of the VIRTIS exposure and a reorientation of the spacecraft close to the equator. Similarly, the footprint of the SPICAV slit moved at a nearly constant longitude, with a small deviation
Figure 1. Spatial coverage of SPICAV (in gray) and VIRTIS-M (in black) nadir observations of the Venus nitric oxide and O2 (1D) airglow during Venus Express orbit 243. 4 of 10
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Figure 2. Examples of concurrent observations of the latitudinal distribution of nightglow intensities at nadir by VIRTIS-M and SPICAV as a function of latitude: (a) orbit 243, (b) orbit 592, (c) orbit 342, and (d) orbit 459. Note the different brightness scales used for the NO and the O2 airglow emission rates. The brightness at 1.27 mm has been corrected for backscattered emission and both emission rates have been corrected for the view angle.
at low latitudes. Other cases of parallel observations with the two instruments present a similar pattern of spatial coverage. Once the regions of observation overlap have been identified, the brightness information is extracted from the VIRTIS nadir images along the track of the SPICAV slit. Since a VIRTIS image is constructed by combining adjacent pixel lines corresponding to successive positions of the mirror, the observations are not exactly coincident in time. The time difference is usually on the order of a few minutes. In the particular case of orbit 243 which covers a wide range of latitudes, the maximum time delay between measurements of the intensity with the two instruments at any given location varies between 4 and 16 min. [13] Figure 2a shows the latitudinal distribution of the NO and O2 (1Dg) nadir emission rates for this orbit measured at 0035 LT between 0752 and 0806 UT on 20 December 2006. The NO brightness is sampled once per second and a smoothing function over 10 s has been applied. The O2 signal has been extracted by averaging the intensity in the processed VIRTIS image over an area corresponding to the projection of one of the SPICAV spatial bins, as was shown
in Figure 1. The region of parallel observations extends from close to the equator up to nearly 50°N. In this case, two different regions are identified. The first one, southward of 25°N presents three successive peaks in the NO intensity with brightness ranging from 1 to 2.7 kR. The O2 airglow latitudinal variation shows structural similarities with the NO variations up to about 14°N, with two peaks nearly coincident with the NO maxima. Further north, no clear increase of the O2 brightness corresponds to NO maximum at 17°N. Northward of the location, no correlation is observed between the two signals. In particular, no feature is observed in the NO intensity at 33°N where the O2 airglow increases by over a factor two. In the lowlatitude region of the two common intensity peaks, the O2/ NO airglow intensity ratio is on the order of 500. A similar case is illustrated by Figure 2b for orbit 592 where the two emissions show a different latitudinal trend equatorward of 24°N, followed by a nearly coincident maximum near 30° reaching 5.6 kR in the NO bands and 1.2 MR at 1.27 mm. The two distributions show little correlation poleward of 35°N. In this example, the average O2/NO intensity ratio
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is again close to 500. An example of uncorrelated structures of the two airglow emissions is illustrated in Figure 2c which shows the latitudinal distribution measured during orbit 342 (29 March 2007) in the premidnight sector (2314 LT). In this case, no correlation is observed between 10°S and 11°S. The O2 airglow exhibits a peak reaching 1 MR on the equator and a decrease on either side of this maximum. A secondary peak is observed at 5°N. The NO airglow presents an equatorial dip with larger intensities up to 3.2 kR at 9°N. In this example, the brightness ratio of the two emissions is on the order of 800. Figure 2d is another example showing no correlation between the NO and the O2 emissions observed during orbit 459. It extends from 20°S to 38°N, close to local midnight (2343 LT). Following correction for the thermal emission component, the O2 (1D) emission rate is very weak southward of 10°N and hardly distinguished from the noise level. It continuously increases toward middle northern latitudes and nearly reaches the 1.2 MR level at the end of the observation sequence. Interestingly, the NO airglow shows a bright maximum of 2.2 kR near 2°N, a region where the O2 emission is very weak. Inversely, the O2 airglow increases poleward of 30°N, which corresponds to a region where the NO airglow drops with increasing latitude. [14] Concurrent sequences of the two airglow emissions have been collected during 14 Venus Express orbits. None of them shows latitudinal distribution of the two emissions which are correlated over the full observation sequence. Instead, the two features are either totally uncorrelated as in Figures 2c and 2d or they exhibit some correlation between the locations of the intensity peaks over a limited latitudinal extent as in Figures 2a and 2b. Table 1 summarizes the dates, times and locations of the parallel airglow observations and the correlation coefficients derived from each orbital sequence. From the examples in Figure 2 and Table 1, we conclude that no large-scale correlation is generally observed in the latitudinal distribution of the vertical emission rate of the NO and O2 airglow. Some of the brightness enhancements are colocated over a restricted region, such as on orbit 343, but the two latitudinal distributions may also show quite a different morphology over regions exceeding 50 degrees of latitude. We now examine concurrent limb observations of the two emissions to verify if the same conclusion holds and increase the sample size of parallel SPICAV-VIRTIS observations.
3. Parallel Limb Observations [15] The observations used for this study were collected in the grazing (tangential) limb mode, where the line of sight is at some angle to the orbital plane and moves in such a way to maximize the time spent in the atmosphere [Titov et al., 2006; Bertaux et al., 2007b]. The SPICAV line of sight scans a range of altitudes, generally between 70 km and 400 km and each second a full UV spectrum is obtained. In this mode, the line of sight crosses the dark limb several times during the ascending portion of the VEX orbit. Therefore, SPICAV supplied several sets of two (one for ingress, one for egress) altitude scans of five altitude profiles at each orbit [Ge´rard et al., 2008a]. During these limb observations, the bin parameter varies between 2 and 32. The apparent altitude of the emission peak and its
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Figure 3. Sketch illustrating the methodology to generate the data point for the study of the O2 and NO concurrent limb observations shown in Figures 5 and 6. The projection of the SPICAV field of view in the atmosphere is represented by the yellow dotted line. The green dotted line shows the parallel line shifted downward by a distance Dh along which the values of the O2 intensity is extracted to trace emission limb profiles such as illustrated in Figure 4.
brightness depend on the value of the field of view projected on the limb. As was discussed by Ge´rard et al. [2008a], the SPICAV field of view projected on the limb intercepts a vertical region whose size depends on the spacecraft limb distance, the orientation of the slit and the bin parameter. It varies from 3 to 27 km, with a mean value of 14.9 km. This effect is accounted for by smoothing by the field of view of an emission layer having a vertical Chapman profile integrated along the line of sight. Each Chapman profile is constrained to show the same topside scale height as the observation. SPICAV data points have been corrected for this smoothing effect by setting the peak intensity and altitude to the values they would have had if the limb profiles had been observed with a negligibly small field of view. The 0.25 mrad pixel size of the VIRTIS-M detector projected on Venus limb corresponds to a spatial resolution of 1.9 km for a spacecraft distance of 7500 km, a typical value for a VIRTIS observation at 40°N. Analysis of the spectral cubes at the limb has indicated that the contribution of thermal radiation from the lower atmosphere is very small in the vicinity of 1.27 mm for altitudes of the tangent point above 85 km and corrections are negligible above 95 km [Piccioni et al., 2009a]. [16] The possible correlation of the altitude and peak intensity of the NO and O2 airglows at the limb has been investigated on a statistical basis. Periods when both SPICAV and VIRTIS-M were observing the same limb region have been identified and concurrent limb profiles of the two airglows layers have been extracted. The methodology consists in first determining the limb profile of the NO airglow distribution during ingress and egress of the SPICAV line of sight in the lower thermosphere and upper mesosphere. Figure 3 illustrates the example of a profile measured for a line of sight ingress. In this sketch, the emission peaks of the two airglow layers are separated by a vertical distance taken to be equal to 15 km on the basis of the NO and O2 limb statistics. Once the NO emission profile is constructed, the corresponding VIRTIS limb image is scanned and pixels with a tangent point altitude located at an altitude Dh below each successive SPICAV data points are extracted from the VIRTIS cube in the 1.27 mm channel.
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Figure 4. Example of concurrent limb profiles of the nitric oxide and O2 infrared airglow measured during Venus Express orbit 323.
In this way, an O2 limb profile is extracted from the VIRTIS-M cube so that the effects of horizontal inhomogeneity of the airglow is minimized. The time separating the acquisition of a SPICAV data point and the underlying O2 intensity is, at most, of a few minutes. In this way, quasisimultaneous profiles of the two emissions are obtained for every ingress or egress when the two instruments were simultaneously operating. Figure 4 shows a example obtained for orbit 323 (10 March 2007) between 0035 and 0047 UT at 2258 LT for an airglow layer separation of Dh = 15 km. The NO limb profile shows a peak of 21 kR at 111 km. The O2 profile reaches 69 MR at 99.8 km with a fast intensity drop above the peak. Statistical error bars are indicated on the VIRTIS and SPICAV data points. A total of 249 such parallel limb profiles have been obtained between 17 January 2007 and 9 January 2008. For each profile, the altitude and brightness of the emission peaks are determined and included in the database. The results of this statistical study are summarized in Figures 5 and 6. [17] The altitudes of the simultaneously observed NO and O2 emission profiles are shown in Figure 5. The scatterplot indicates that the intensity of the two airglow layers is not correlated, as confirmed by the very low value of the correlation coefficient R = 0.05. The dashed line indicates equal altitudes for the NO and O2 emission peaks. This plot also confirms that the O2 emission peak is in most circumstances located below the NO layer. However, the distance separating the two emission peaks varies from nearly zero to as much as 28 km, with an average of 15 km. A detailed analysis of the four data points where the O2 peak is above the NO peak has been looked at in detail. These are specific cases when the NO airglow intensity shows a considerable latitudinal gradient as evidenced by the different peak altitudes obtained with the different SPICAV spatial bins. Since the method we use to extract the O2 limb profiles from the VIRTIS images is such that the NO and O2 profiles do not exactly correspond to the same observed volume, it is possible that the actual O2 emission peaks are not really located above the NO emission. The star indicates that the average altitude of the NO airglow peak for this data set is 113 km and 96.4 km for O2(1D). The large scatter in the distance between the two airglows layers probably reflects
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Figure 5. Altitude distribution of the emission peaks of the NO and O2 (1D) airglow layers measured quasisimultaneously at the limb. The observations corresponding to each point have been collected as illustrated in Figures 3 and 4 (see text). The dashed line indicates equal altitudes for the two airglow emission peaks. The correlation coefficient is 0.05, indicating the lack of covariation of the altitude of the two emissions. The full square indicates the mean value of the peak altitude of the two emissions.
the widely changing dynamical regime prevailing in the transition region between the upper mesosphere and the lower thermosphere. A similar result is obtained when comparing the brightness of the emission peaks in parallel observations shown in Figure 6. For clarity, the NO and O2 brightness has been plotted on a logarithmic scale since they vary by over a wide range of values. As it was found for the peak altitude, this plot indicates that the limb brightness of the two emissions is not significantly correlated (R = 0.29). The O2/NO intensity ratio varies by nearly 3 orders of magnitude from 55 to 40,000. The black square indicates
Figure 6. Brightness distribution of the emission peaks of the NO and O2 (1D) airglow layers measured quasisimultaneously at the limb. The observations corresponding to each point have been determined as illustrated in Figures 3 and 4 (see text). The correlation coefficient is 0.29, indicating the lack of covariation of the brightness of the two emissions. The full square indicates the mean value of the peak intensity of the two emissions.
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that the average limb brightness of the NO and O2 is 45 kR and 35 MR, respectively, in this data set, in good agreement with the statistical values of 32 kR for NO [Ge´rard et al., 2008a] and 29 MR for O2(1D) (A. Saglam et al., Characteristics of the OH infrared nightglow in the Venus mesosphere and correlation with the O2 (1D) emission at 1.27 mm based on VIRTIS limb observations, submitted to Advances in Space Research, 2009). The sensitivity of the results to the value of the distance Dh separating the two emission peaks has been tested by varying it by ±5 km. No significant difference was found in the results. It is thus concluded that neither the altitude of the emissions peaks nor the peak intensities at a given location in the Venus nightside atmosphere are correlated. This result is in full agreement with the independent conclusion derived from the nadir observations reported in section 2. We now examine possible explanations for these differences in the section 4.
4. Discussion [18] The observations collected with the SPICAV spectrograph and the VIRTIS-M spectral imager provide evidence that the molecular oxygen and the nitric airglow emissions are only weakly correlated. This conclusion is based on three different sets of observations. First, the statistical location of the NO and the O2(1D) regions of enhanced emission are not coincident. This was demonstrated by the difference between the O2 spot centered on the equator at midnight [Ge´rard et al., 2008b; Piccioni et al., 2009a] and the NO maximum which is shifted by about two hours toward dawn and southward of the equator [Stewart et al., 1980; Bougher et al., 1990]. Second, nadir quasi-simultaneous observations of the two emissions reported in section 2 of this study demonstrate that the distribution of the intensity along latitudinal cuts exhibits significant differences even though the two emissions may show similarities over a limited range. Third, limb observations indicate that neither the brightness nor the altitude of the emission peak covary in the northern hemisphere. The first aspect requires additional studies and modeling. The shift of the statistical region of bright NO emission from the antisolar point [Stewart et al., 1980; Bougher et al., 1990; Bougher and Borucki, 1994] is an indication that the dawnward superotation observed at the cloud level persists in the upper hemisphere in such a way that the subsidence region of the global thermospheric circulation is statistically displaced by 2 h. This result needs further SPICAV observations to confirm that the shift is still observed in a period of low solar activity conditions. If the difference between the locations of the NO and O2 bright regions is still observed during the Venus Express era, the picture of the vertical wind structure in the upper mesosphere-lower thermosphere transition region has to be revised accordingly. [19] The decoupling between the characteristics of the two emissions at a given location and time raises a different question. At first glance, the results reported in this study contradict the concept of a global subsolar-to-antisolar circulation carrying the O and N atoms from their dayside source region to the nightside location where they recombine to produce the NO and O2 (1D) airglow emissions. In this view, the region of subsidence of the two species would
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Figure 7. Sketch illustrating the role of horizontal wind in the mesosphere-thermosphere transition region as a source of spatial decorrelation between a bright spot of NO airglow and a region of enhanced O2 (1D) nightglow. The region initially enriched in oxygen atoms has traveled a horizontal distance DL by the time the blob of O-rich gas reaches the altitude of the O2 nightglow layer located Dz km below the O2 emission peak.
be nearly coincident in the vicinity of the antisolar point. We note however that, at any given time, airglow images indicate that the location of the regions of bright emissions may considerably vary as was shown for NO [Stewart et al., 1980; Bougher et al., 1990] and O2 (1D) [Hueso et al., 2008; Piccioni et al., 2009a]. This important feature shows that the location of the subsidence of the SS-AS circulation is very variable, presumably as a consequence of an intrinsic variability of the circulation (possibly caused by the drag from gravity waves) and horizontal winds. As mentioned before, apparent wind velocities locally as large as about 100 ms1 have been deduced by Hueso et al. [2008] from the displacement of O2 airglow features between successive VIRTIS-M images. [20] We now take a close look at the role of horizontal winds in the decorrelation between the two airglow emissions. The two airglow layers are separated in altitude typically by 16 km. Although this distance is limited, the time required for vertical down transport of atoms from the NO layer to the O2 emission peak is fairly long, leaving the possibility that a region richer in oxygen at 113 km has moved over considerable horizontal distances by the time it reaches the 97 km level. This situation is illustrated by the sketch in Figure 7 where an air parcel of enhanced NO emission (in white-blue) is carried over an horizontal distance DL by the time it has reached the altitude of maximum O2 recombination Dh km below the NO layer. The effect of this horizontal transport in a situation of nonuniform downward flux above the NO airglow layer is potentially able to explain the observed decorrelation. We first need to estimate the downward velocity of an air parcel to traveling the distance Dh separating the two airglow layers. The vertical transport velocity w has not been directly measured but may be estimated using models and circumstantial evidence. We examine three possible ways to derive typical vertical velocity values. First, in the one-dimensional chemical
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transport model used by Cox et al. [2008] and Ge´rard et al. [2008a], vertical transport below the homopause is parameterized by an eddy diffusion coefficient K. We adopt a standard value of DZ = 16 km for the distance separating the average altitudes of the NO and O2 emission peaks. On the basis of numerical simulations using the expression K = A/n1/2, we find a velocity on the order of 3 cm s1. A second approach to estimate w is based on results from the three-dimensional model by Bougher et al. [1990] where eddy diffusion is relatively small. In a recent version of this model, the vertical advection velocity between 115 and 95 km is on the order of 15–20 cm s1 near the antisolar point (S. W. Bougher, private communication, 2009). Finally, an estimate of the downflow velocity was also given by Bailey et al. [2008] for a nightside region with a 20 K temperature enhancements observed over a 2-day period. They obtained a vertical flow velocity of 20 cm s1 for such a region of airglow brightening associated with enhanced vertical transport. The corresponding transit time for atoms flowing from the NO to the O2 airglow peak ranges between 22 h and 6 Earth days. [21] The horizontal wind velocity v at the level of the O2 airglow layer is very variable as was recently summarized by Lellouch et al. [2008]. Winds velocities, derived from CO millimeter observations, are typically on the order of 30– 50 m s1 at 93 km and 90– 120 m s1 near 102 km. Hueso et al. [2008] derived values of the effective wind velocity at 97 km of a few tens of m/s from the motions of bright spots of O2 airglow. Horizontal winds calculated with the 3-D model values are on the same order in the region separating the two emission layers. A crude estimate of the horizontal distance DL crossed by moving air parcel is then given by DL = v/w DZ. Adopting a value of v = 100 m s1 as an upper limit, we find values on the order of 53,000 km if vertical transport is parameterized by eddy diffusion or 8000 km if the estimate for w from Bailey et al. adopted. Using v = 10 m s1 as a wind velocity probably closer to the average, we estimate a typical horizontal transport range of 800 to 5300 to km. Actually, the situation is further complicated by the occurrence of chemical reactions and collisional quenching which limit the chemical lifetimes of the ground state O atoms and the excited O2(1D) molecules, respectively. In any case, it appears that the downward flow is much slower than the horizontal transport and atoms may travel considerable horizontal distances during their transit from 113 to 97 km. It is therefore a direct explanation of the lack of correlation between the two emissions observed concurrently by SPICAV and VIRTIS-M. The very fact that the two emissions exhibit significant differences in their horizontal distribution strongly argues for the presence of strong horizontal winds in the thermosphere-mesosphere transition region. Other factors such as the different altitude where the N and the O atoms are formed on the dayside may play an additional role in the lack of correlation since the streamlines followed during their transport to the nightside are slightly different. [22] On the basis of these arguments, it may now appear more difficult to understand the covariation occasionally observed in latitudinal cuts such as in Figures 2a and 2c. We speculate that the two emissions may covary at least under three particular conditions. First, when the distance between the two airglow layers is less than the average 16 km. The
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existence of such conditions is testified by Figure 6, which shows the occurrence of cases when this distance may be reduced to only a few kilometers. Under such circumstances, the vertical downflow time and the horizontal traveled distance may also be reduced accordingly. Second, maps of the horizontal displacements of bright O2 features have indicated that regions with much smaller horizontal velocities, sometimes quasi-null values, have been observed near 97 km. It is possible that these horizontal stagnation regions persist over a sufficiently long period of time, allowing a more vertical downflow of the oxygen atoms. Finally, local enhancements in the brightness of the O2 airglow suggest the presence of strong downflows, possibly associated with an increased vigor of local downward turbulent transport. These conditions may well explain that the two emissions can exhibit some degree of covariation, which is otherwise absent in areas of strong horizontal winds and/or weaker vertical velocity.
5. Conclusion [23] The O2(1D) and the NO d and g bands night airglows of Venus have been observed concurrently for the first time by two instruments on board the Venus Express spacecraft. These observations include both nadir and limb viewing geometries. During nadir observations, the O2 (a1D) emission intensity has been extracted from VIRTIS images observing the same locations as the footprint of the SPICAV slit in the Venus atmosphere on the same orbit. Occasional positive correlations between the latitudinal distributions of the two emissions have been observed over a limited latitudinal range. However, the overall covariability of the two airglows is low, as shown by some of the results illustrated in this study and by the globally low correlation coefficients obtained between the two sets of observational sequences. A similar conclusion is reached from the statistical comparison of the altitude and peak brightness of the two emissions simultaneously observed in the grazing limb geometry. The very low correlation coefficients of both the altitude and the intensity of the two airglow layers indicate that the transport of O and N atoms considerably deviates from a steady state vertical flow, even in the region close to the antisolar point. We suggest that the airglow decorrelation frequently observed is a consequence of the transport of the downward moving air mass by strong horizontal winds in the transition region that dynamically decouples the airglow in the two layers in a given vertical column. Our simple calculation of horizontal transport indicates that the atoms may travel considerable distances during the transit time of their vertical transport by vertical advection and eddy diffusion. Occasional correlations of the latitudinal intensity variations may be associated with situations when the distance between airglow layers is small or when the ratio of the horizontal and vertical transport velocity components is reduced. Simulations with a two- or threedimensional chemical transport model are needed to assess this scenario and quantify this effect. [24] Acknowledgments. We gratefully thank all members of the ESA Venus Express project and of the VIRTIS ( http://servirtis.obspm.fr/Venus_ Express/VIRTIS_Team.html) and SPICAV scientific and technical teams. J.C.G. acknowledges funding from the Belgian Fund for Scientific Research (FNRS). A. Saglam, C. Cox, and L. Soret were supported by
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the PRODEX program managed by the European Space Agency with the help of the Belgian Federal Space Science Policy Office. This work was funded by Agenzia Spaziale Italiana and the Centre National d’Etudes Spatiales.
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SPICAV on board Venus Express, J. Geophys. Res., 113, E00B03, doi:10.1029/2008JE003078. Ge´rard, J.-C., A. Saglam, G. Piccioni, P. Drossart, C. Cox, S. Erard, R. Hueso, and A. Sa´nchez-Lavega (2008b), The distribution of the O2 infrared nightglow observed with VIRTIS on board Venus Express, Geophys. Res. Lett., 35, L02207, doi:10.1029/2007GL032021. Ge´rard, J.-C., A. Saglam, G. Piccioni, P. Drossart, F. Montmessin, and J.-L. Bertaux (2009), Atomic oxygen distribution in the Venus mesosphere from observations of O2 infrared airglow by VIRTIS-Venus Express, Icarus, 199, 264 – 272, doi:10.1016/j.icarus.2008.09.016. Hueso, R., A. Sa´nchez-Lavega, G. Piccioni, P. Drossart, J.-C. Ge´rard, I. Khatuntsev, and L. Zasova (2008), Morphology and dynamics of Venus oxygen airglow, J. Geophys. Res., 113, E00B02, doi:10.1029/ 2008JE003081. Krasnopolsky, V. A. (1986), Oxygen emissions in the night airglow of the Earth, Venus, and Mars, Planet. Space Sci., 34, 511 – 518, doi:10.1016/ 0032-0633(86)90089-9. Lellouch, E., T. Clancy, D. Crisp, A. Kliore, D. Titov, and S. W. Bougher (1997), Monitoring of mesospheric structure and dynamics, in Venus II: Geology, Geophysics, Atmosphere, and Solar Wind Environment, edited by S. W. Bougher et al., pp. 295 – 324, Univ. of Ariz. Press, Tucson. Lellouch, E., G. Paubert, R. Morenoaand, and A. Moullet (2008), Monitoring Venus’ mesospheric winds in support of Venus Express: IRAM 30-m and APEX observations, Planet. Space Sci., 56, 1355 – 1367, doi:10.1016/j.pss.2008.06.010. Miller, H. C., J. E. McCord, J. Choy, and G. D. Hager (2001), Measurement of the radiative lifetime of O2 (a1Dg) using cavity ring down spectroscopy, J. Quant. Spectrosc. Radiat. Transfer, 69, 305 – 325. Ohtsuki, S., N. Iwagami, H. Sagawa, M. Ueno, Y. Kasaba, T. Imamura, and E. Nishihara (2008), Imaging spectroscopy of the Venus 1.27-mm, O2 airglow with ground-based telescopes, Adv. Space Res., 41, 1375 – 1380, doi:10.1016/j.asr.2007.10.014. Piccioni, G., L. Zasova, A. Migliorini, P. Drossart, A. Shakun, A. Garcı´a Mun˜oz, F. P. Mills, and A. Cardesin-Moinelo (2009a), Near-IR oxygen nightglow observed by VIRTIS in the Venus upper atmosphere, J. Geophys. Res., 114, E00B38, doi:10.1029/2008JE003133. Piccioni, G., et al. (2009b), The Visible and Infrared Thermal Imaging Spectrometer, Eur. Space Agency Spec. Publ., ESA SP-1295, in press. Saglam, A., J.-C. Ge´rard, L. Soret, G. Piccioni, and P. Drossart (2009), Characteristics of the OH infrared nightglow in the Venus mesosphere and correlation with the O2 (1D) emission at 1.27 mm based on VIRTIS limb observations, submitted to Advances in Space Research. Stewart, A. I., and C. A. Barth (1979), Ultraviolet night airglow of Venus, Science, 205, 59 – 62, doi:10.1126/science.205.4401.59. Stewart, A. I. F., J.-C. Ge´rard, D. W. Rusch, and S. W. Bougher (1980), Morphology of the Venus ultraviolet night airglow, J. Geophys. Res., 85, 7861 – 7870, doi:10.1029/JA085iA13p07861. Svedhem, H., et al. (2007), Venus Express: The first European mission to Venus, Planet. Space Sci., 55, 1636 – 1652, doi:10.1016/j.pss.2007. 01.013. Titov, D. V., et al. (2006), Venus Express science planning, Planet. Space Sci., 54, 1279 – 1297, doi:10.1016/j.pss.2006.04.017.
J.-L. Bertaux, LATMOS, Re´duit de Verrie`res, BP3, Route des Gaˆtines, F-91371 Verrie`res-le-Buisson, France. (
[email protected]) C. Cox, J.-C. Ge´rard, A. Saglam, and L. Soret, Laboratoire de Physique Atmosphe´rique et Plane´taire, Universite´ de Lie`ge, B-4000 Lie`ge, Belgium. (
[email protected];
[email protected];
[email protected]; lauriane.
[email protected]) P. Drossart, LESIA, Observatoire de Paris, 5 Place Jules Janssen, F-92195 Meudon, France. (
[email protected]) G. Piccioni, INAF, IASF, Via del Fosso del Cavaliere 100, I-00133 Rome, Italy. (
[email protected])
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Venus surface thermal emission at 1 mm in VIRTIS imaging observations: Evidence for variation of crust and mantle differentiation conditions N. Mueller,1 J. Helbert,2 G. L. Hashimoto,3 C. C. C. Tsang,4 S. Erard,5 G. Piccioni,6 and P. Drossart5 Received 15 February 2008; revised 10 July 2008; accepted 20 August 2008; published 10 December 2008.
[1] The Venus Express spacecraft images the nightside thermal emissions using
the Visible and Infrared Thermal Imaging Spectrometer (VIRTIS). At 1.02, 1.10, and 1.18 mm, thermal emission from the surface is observed. The signal is attenuated by scattering and absorption in the dense atmosphere. The measured flux at the top of the atmosphere is positively correlated with surface temperature and surface emissivity. The surface temperature of Venus is relatively well constrained as being mainly a function of altitude with a gradient lesser or equal to the adiabatic lapse rate. This study examines the correlation of VIRTIS images showing a signal of the surface at 1.02 mm with viewing geometry, stray sunlight, cloud opacity, and topography and applies semiempirical relations to remove their influence. The remaining contrast can be either ascribed to surface emissivity or unexpected temperature variations. Temperature variations due to active volcanism are unlikely to be persistent over the time of observations; therefore, the mosaic of all processed images is here interpreted in terms of surface emissivity variation. The emissivity variation found is correlated with geomorphological features established from Magellan synthetic aperture radar images. It is generally lower at tessera terrain. Some, but not all, volcanic edifices show increased emissivity. Large lava flows in the Lada terra-Lavinia planitia region also show an increased thermal emission. This might indicate a more felsic surface composition of tessera highlands and large-scale extrusive volcanism of ultramafic composition. Citation: Mueller, N., J. Helbert, G. L. Hashimoto, C. C. C. Tsang, S. Erard, G. Piccioni, and P. Drossart (2008), Venus surface thermal emission at 1 mm in VIRTIS imaging observations: Evidence for variation of crust and mantle differentiation conditions, J. Geophys. Res., 113, E00B17, doi:10.1029/2008JE003118.
1. Motivation [2] The search for exoplanets yields an ever increasing number of known planets. With further improvements in observation techniques it will hopefully soon be possible to detect Earth sized planets orbiting their stars at distances where liquid surface water is stable. The special interest in planets with these attributes is of course motivated by the desire to find habitable planets or even life. Size and orbital distance comparable to Earth is by no means an ideal indicator of habitability as the Earth’s neighboring planets demonstrate. Venus with eight tenths of Earth’s mass and 1 Joint Planetary Interior Physics Research Group of the University Muenster and DLR Berlin, Berlin, Germany. 2 Institute of Planetary Research, DLR, Berlin, Germany. 3 Laboratory for Earth and Planetary Atmospheric Science, Department of Earth and Planetary Sciences, Kobe University, Kobe, Japan. 4 Atmospheric, Oceanic, and Planetary Physics, Department of Physics, Clarendon Laboratory, University of Oxford, Oxford, UK. 5 LESIA, Observatoire de Paris, Paris, France. 6 IASF-INAF, Rome, Italy.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003118$09.00
nearly three fourths of Earth’s orbital distance is quite inhospitable. The surface of Venus is extremely hot owing to greenhouse climate imposed by its dense CO2 atmosphere and sulfuric acid clouds. Surface and atmosphere are also dry compared to Earth [Taylor, 2006]. [3] Comprehensive understanding of why Venus so utterly failed to become remotely habitable has not yet been achieved [Crisp et al., 2002]. One missing piece for the reconstruction of the evolution of Venus is global knowledge on the composition of the crust. The surface has been extensively mapped at radar wavelengths by the Magellan mission [Pettengill et al., 1991] which further revealed a unique geology [e.g., Saunders et al., 1991; Head et al., 1992; Solomon et al., 1992; Schaber et al., 1992; Phillips et al., 1992; Stofan et al., 1992, 1997; Hansen et al., 1997; Smrekar et al., 1997; Grimm and Hess, 1997; Basilevsky et al., 1997]. Radar is however not very sensitive to mineral composition. The discovery of spectral windows in the atmosphere (Figure 1) that allow measurable amounts of near-infrared thermal emission from the surface to escape [Carlson et al., 1991] presents the opportunity for limited remote sensing of surface mineralogy [Carlson et al., 1993a; Lecacheux et al., 1993; Baines
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Figure 1. Modeled spectrum of the near-infrared (NIR) windows at VIRTIS-H spectral resolution [Drossart et al., 2007]. Model is described by Tsang et al. [2008]. The different windows feature different amounts of emission originating from the surface: at 1.02 mm more than 95% comes from the surface, and at 1.27 and 1.31 mm a negligible part of the radiation comes from the surface [Meadows and Crisp, 1996]. et al., 2000; Moroz, 2002; Hashimoto and Sugita, 2003; Hashimoto et al., 2008; Arnold et al., 2008; A. T. Basilevsky et al., Geologic interpretation of the near-infrared images of the surface taken by the Venus Monitoring Camera, Venus Express, submitted to Journal of Geophysical Research, 2008]. [4] The window most suited for this purpose is at 1 mm which coincides with a absorbtion band of FeO, one of the primary compounds of mafic minerals. Abundance of refractory mafic minerals gives evidence of the conditions under which the ultramafic mantle material differentiated and formed the crust. It is possible to distinguish three different types of crustal differentiation [Taylor, 1989]. Primary differentiation is the formation of floating crust in a fully molten mantle or magma ocean. Evidence for this is found in the lunar highland anorthosite rocks returned by the Apollo missions [Taylor, 1974]. Secondary differentiation occurs when the convecting mantle material partially melts, the liquid phase moves upward and forms a basaltic crust either by volcanism or plutonic intrusion, e.g., at midocean ridges on Earth. Tertiary differentiation requires recycling of basaltic crust and water into the mantle. Water lowers the liquidus temperature of secondary crust which creates partial melts with a felsic composition that generated the bulk of the continental crust on Earth [Taylor and Campbell, 1983]. The greater part of the venusian crust is thought to be formed by secondary differentiation [Head et al., 1994; Grimm and Hess, 1997]. Arguments for this are lava viscosity inferred from morphology, in situ measurements and the unimodal hypsometry. Most in situ measurements indicate a mafic composition comparable to common terrestrial basalts [Surkov and Barsukov, 1985; Surkov et al., 1986, 1987]. Two out of seven, Venera 8 and Venera 13, indicate a more complex differentiation process that involves enrichment in abundance of alkalines [Kargel et al., 1993]. Abundance of alkalines is not reflected in the felsic-mafic description.
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[5] Crustal differentiation is counteracted by thermal buoyancy, which on Earth leads to remixing of crustal material into the mantle by oceanic plate subduction. It is not very well constrained whether crustal recycling takes or took place on Venus [Grimm and Hess, 1997]. Crater frequency gives strong evidence for relative recent resurfacing [Phillips et al., 1992; Schaber et al., 1992; Strom et al., 1994]. The crust mapped by radar appears young because extrusive volcanism and tectonic deformation have erased the early crater history. The processes responsible for the resurfacing are not fully understood, and the fact of a seemingly young crust does not constrain crustal recycling. For instance, the concept of episodic plate tectonics [Turcotte, 1993] involves recycling of most of the crust and simultaneous creation of new crust. In other models considering vertical rather than horizontal accretion [e.g., Parmentier and Hess, 1992; Head et al., 1994], recycling takes place only if crustal growths exceeds a thickness of 60 to 80 km and crustal basalt minerals undergo a phase transition to denser eclogite [Spohn, 1991]. Mapping of geochemically old crust (i.e., granite or anorthosite) might constrain rate of crustal formation and recycling. [6] Another product of mantle differentiation is the silica and FeO depleted material remaining in the mantle after the liquid basaltic phase migrated upward. Parmentier and Hess [1992] assume that this residual material forms a compositionally buoyant layer below the crust. On Earth this reservoir is thought to be small owing to lithosphere subduction. On Venus plate tectonics is not unambiguously identifiable [Kaula and Phillips, 1981; Solomon et al., 1992], and the lithosphere is assumed at present to form a stagnant lid [Solomatov and Moresi, 1996]. The reservoir of depleted mantle material on Venus is therefore likely greater than on Earth and steadily increasing during secondary differentiation as it is not remixed into the mantle by subduction. Cooling of this conducting layer leads to negative net buoyancy and foundering of the layer drives episodic tectonic and volcanic resurfacing. In the context of this resurfacing model, Head et al. [1994] predict composition of lava to be of ultramafic, e.g., picritic or komatiitic composition during the periods of quiescence as partial melting in mantle diapirs occurs in deeper regions under different conditions. [7] Regardless of the stagnant lithosphere the surface is strongly influenced by endogenous processes. Mantle diapirs supposedly lift the volcanic dome shaped highlands through dynamic support and lithosphere thinning and imprint coronae structures [Smrekar et al., 1997; Stofan et al., 1997]. The wide range of topographic expressions of coronae can be explained by mantle diapirism triggering a gravitational instability and delamination of lithosphere [Smrekar and Stofan, 1997]. Lithosphere delamination can also account for the wide range of inferred lava viscosities requiring a diversity of melt compositions [Elkins-Tanton et al., 2007]. Their modeling of the temperature-pressure conditions in and around a diapir of delaminating lithosphere demonstrates the possibility of generation of melt varying in SiO 2 abundance from 60% (intermediary between felsic and mafic) to 40% (ultramafic). [8] The process of differentiation and recycling of crustal material is relevant for the thermal evolution of the interior of the planet. One side is the effective cooling of the interior
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by recycling of cool crustal material. The other side is the distribution of radioactive heat sources throughout the planet. Important radioactive heat sources such as 40K, 140 Th, and 239U are incompatible, during partial melting these isotopes are enriched in the liquid, more felsic phase. Thus, on Earth the oceanic basaltic and especially the continental granitic crust contain higher abundances of radioactive elements than the mantle. This concentration of heat sources toward the top decreases the heat flux in the deep interior of the planet and also influences amount of magmatism by depletion of mantle heat sources [Spohn, 1991]. [9] The convective heat flux in the outer core driven by overall cooling of the planet is the energy source for the magnetic field of the Earth and presumably Mercury [Stevenson et al., 1983]. The geodynamo process is currently not active on Venus but it is conceivable that it was in the past when distribution of the radioactive elements or the cooling of the planetary interior by lithosphere recycling was different. An intrinsic magnetic field would greatly influence the process of atmospheric erosion and allow Venus to retain its primordial water longer [Donahue and Russell, 1997]. In such an primordial climate and tectonic regime tertiary differentiation might have occurred and led to granitic cratons still existing in the Venusian highlands [Taylor and Campbell, 1983]. Hashimoto and Sugita [2003] concluded that it is possible to distinguish mafic basalt from felsic granite or rhyolite from orbit when given a sufficient but not impossible accuracy of measurement of thermal emission on the nightside of Venus. [10] The aim of this study is to analyze VIRTIS images to find whether there is a spatial signal that can be plausibly attributed to content of mafic minerals of the topmost crust in the context of surface morphology from Magellan synthetic aperture radar (SAR) images.
2. Background and Theory 2.1. Atmosphere and Surface Temperature [11] The thermal emission is dominated by surface temperature T as it is at any wavelength l the product of surface emissivity el and blackbody spectral radiance given by the Planck function Bl(T). At the wavelengths and temperatures relevant here (1 mm and 735 K) the Planck function varies strongly with temperature. To retrieve surface emissivity from remote sensing of the thermal emission, the physical surface temperature must be well known. On Venus atmospheric temperature is well constrained owing to the insulating greenhouse climate and measurements at different local times and locations by the Venera, Pioneer Venus, and Vega descent probes. The surface temperature is thought to be in thermal equilibrium with the atmosphere owing to high heat capacity of the atmosphere and little diurnal variations [Lecacheux et al., 1993]. [12] Stone [1975] estimated the different time scales of radiative heating and cooling and convective heat transport of the atmosphere of Venus and concluded that temperature was governed by the latter. Change of temperature by solar heating and cooling during the night is effectively distributed through the whole atmosphere by advection. Average diurnal and meridional temperature variations in the lower atmosphere were estimated to be less than 0.1 K.
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[13] The in situ measurements of Venera 9 to 11 and the four Pioneer Venus descent probes reviewed by Seiff [1983] agree with each other within approximately 12 K at the same pressure level in the atmosphere below the clouds. The temperature increases toward the surface close to the adiabatic gradient. Diurnal variations are less than 1 K and latitudinal variations are less than 5 K in the data from the Pioneer Venus probes. [14] Observations combined with atmospheric models lead to the Venus International Reference Atmosphere (VIRA). The model temperature structure is presented by Seiff et al. [1985]. The atmosphere in contact with the surface close to the mean planetary radius has a temperature of 735.3 K at 92.1 bar, adiabatical lapse rate is 8.06 K km1, no diurnal or latitudinal variations are included in the model atmosphere. [15] Little is known however about the planetary boundary layer of the atmosphere since Venera probes were designed to travel through the hot atmosphere at great speed to maximize lifetime on the surface and the vertical temperature sampling rate is low. The higher-resolution measurements of the Pioneer Venus probes ceased at 12 km height. The Vega 2 temperature profile is the only measurement in the lower part of the atmosphere comparable in resolution to Pioneer Venus measurements. It shows a stratified atmosphere with dynamically unstable superadiabatic lapse rates in the lowest 6.5 km [Seiff, 1987]. In the lowest 1.5 km the lapse rates vary between 1.5 and 10 K km1. These measurements are not verified and no explanation for them has been found yet [Crisp and Titov, 1997]. Theoretical consideration by Gierasch et al. [1997] of the planetary boundary, where vertical convective motion is inhibited by the surface of the planet, gives an upper estimate of a diurnal temperature oscillation of 6 K corresponding to a 170 m thin thermal boundary layer. [16] In conclusion, the small variation of temperature in the lowest layer of the atmosphere is well constrained but the temperature structure is not. The single high-resolution in situ measurement points toward possible superadiabatic lapse rates near the surface but is on average 8 K km1. Temperatures retrieved from observation of the thermal surface emission at 1.18 mm via radiative transfer modeling suggest a subadiabatic lapse rate of less than 7.5 K km1 [Meadows and Crisp, 1996]. Measurement variations and theoretical diurnal variations near the surface are both less than 10 K at the same pressure level. The atmospheric boundary layer is probably not very relevant for the nightside owing to the long duration of nights on Venus (58 Earth days). No significant near-infrared (NIR) thermal emission variation with local time or latitude has been reported by Lecacheux et al. [1993], Meadows and Crisp [1996], or Hashimoto et al. [2008]. 2.2. Radiative Transfer in the Atmosphere of Venus [17] The atmosphere of Venus consists mostly of CO2 (96.5 wt %) and N2 (3.5 wt %) [von Zahn et al., 1983] with a pressure of 93.2 bar at the zero altitude mean planetary radius (MPR) [Seiff et al., 1985]. From approximately 50 to 70 km above MPR there are several cloud layers composed mostly of droplets of concentrated sulphuric acid. The clouds have spatially varying number densities and size distributions resulting in optical depths between 30 and
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Figure 2. The altitude where brightness temperature of the lower atmosphere without clouds corresponds to atmospheric temperature gives an approximation for the source region of radiation. Spectrum calculated from the same model as in Figure 1 but without clouds. Isolines of Rayleigh scattering optical depths t r calculated following Hansen and Travis [1974]. Rayleigh scattering is less relevant for the 1.31 mm window than for the 1.02 mm window. Brightness temperature varies little between 1.01 and 1.03 mm in this model with a constant continuum absorption coefficient. Compare with spectral dependance of brightness temperature in Figure 8. 50 at visible wavelengths, for a review see Ragent et al. [1985]. A three modal size distribution with modal radii of 0.3, 1.0, and 3.7 mm is well established from in situ and remote observations [Hansen and Hovenier, 1974; Marov et al., 1980; Knollenberg and Hunten, 1980; Carlson et al., 1993b; Grinspoon et al., 1993], but it is unclear whether this is a comprehensive description of all cloud particles. In particular the nature of the UV absorber leaving distinct cloud shaped marks on dayside images is not known [Esposito et al., 1997]. Also, for mode 1 and 3 particles, compositions different from H2SO4 solutions and including solid crystals have been proposed [Ragent et al., 1985; Esposito et al., 1997]. In clouds models usually a mode 20 is included that features a slightly larger modal radius and size variation than mode 2. Above and below the clouds are haze layers of mostly submicrometer particles. Hazes below the clouds have a total optical depth of less than 3 [Ragent et al., 1985] that, however, might have been as high as 10 in Venera 8 near terminator measurements [Moroz, 2002]. Ragent et al. [1985], reviewing the different in situ measurements of hazes, gives no information for altitudes below 10 km. Grieger et al. [2004] reexamined the zenith and nadir radiance measurements of Venera 13 and Venera 14 and found in both a distinct peak of extinction between 1 and 2 km that they concluded to be a near surface cloud deck or haze layer. [18] In conclusion, the atmosphere is optically thick at all visible and infrared wavelengths. Light scattering in the clouds and hazes prohibits direct or indirect surface observation in visible light. Scattering is conservative [Ragent et al., 1985] which allows a small fraction of incident sunlight to reach the surface [see, e.g., Gierasch et al., 1997]. Most
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sunlight is reflected at the cloud layer but Rayleigh scattering in the deep atmosphere is also not negligible [Sagan, 1962; Moroz, 2002]. Sunlight reflected from the surface is not significantly distinguishable against this background down to an altitude of 2 km at 0.65 mm [Moroz, 2002]. Figure 2 shows altitudes of Rayleigh scattering optical thicknesses 0.1 and 1 in the wavelengths studied in here. [19] In the infrared range beyond 2.5 mm gases and H2SO4 cloud particles become strongly absorbing, thermal emission corresponds approximately to the temperatures of the cloud tops. The average brightness temperature of Venus in this range is 220 to 250 K on both dayside and nightside [Pettit and Nicholson, 1955; Moroz et al., 1985] which corresponds to atmospheric temperature at 60 to 75 km height [Seiff et al., 1985]. In the near infrared region there are several windows between CO2 and H2O absorption bands that allow measurable thermal emission from deeper layers to escape. The emissions from these windows were first discovered by Allen and Crawford [1984] at 1.74 mm and 2.30 mm. On the basis of shape, distribution, and revolution period of the bright and dark markings they concluded the observable contrast to be due to the illumination of the variable lower cloud layer by thermal emissions of the hot matter below the clouds. Kamp et al. [1988] found the emitting matter at these wavelengths to be CO2 in the deep atmosphere. Absorption/emission increases with depth owing to higher density and pressure and temperature broadening of absorption bands. On the basis of their calculations they predicted further windows in the 1 mm region that were subsequently observed at 1.10 mm, 1.18 mm, 1.27 mm and 1.31 mm by ground-based observations [Crisp et al., 1991] and during the Galileo flyby [Carlson et al., 1991] at 1.2 mm, 1.01 mm, and 0.8 mm. Two more windows were observed by Cassini VIMS at 0.85 mm and 0.9 mm [Baines et al., 2000]. [20] In addition to the cloud like markings simultaneously seen in all window regions, in those shortward of 1.2 mm the contrast of the thermal emission of cold highlands and hot lowlands is observed [Crisp et al., 1991; Carlson et al., 1993a]. Lecacheux et al. [1993] investigated this contrast in ground-based observations of the 1.02 mm window. Using a radiative transfer model they found the contrast correlated with surface elevation to be consistent with thermal emission of the surface in thermal equilibrium with the atmosphere and less than 10% variation in emissivity. [21] One obstacle in modeling of the windows into the lower atmosphere of Venus is that little or no laboratory data exist on gaseous absorption at comparable pressures and optical paths. In spectral window regions far from absorption bands, absorption coefficients cannot fully be determined from existing absorption line databases and standard line shapes [Taylor et al., 1997]. Parameters describing the continuum absorption are often determined empirically, i.e., chosen such that modeled spectra best fit the observed spectra [see, e.g., Pollack et al., 1993; Meadows and Crisp, 1996; Marcq et al., 2006; Tsang et al., 2008]. [22] The optical properties of the clouds are better understood and consistent with Mie scattering at spheres with the refractive index of 75 wt % solution of H2SO4 [e.g., Hansen and Hovenier, 1974; Grinspoon et al., 1993; Pollack et al., 1993; Carlson et al., 1993b]. On the basis of imaging observations of the windows at 1.74 mm and 2.30 mm with
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Galileo NIMS Carlson et al. [1993b] attributed the cloud contrast not only to varying particle number densities but also to a shift in size or modal distribution. Grinspoon et al. [1993] calculated the scattering properties of cloud particles using Mie theory. Subsequent radiative transfer modeling yielded best results to NIMS data by variation of modes 2’ and 3 in the middle and lower clouds, confirming the findings of Carlson et al. [1993b]. [23] Figure 1 shows a synthetic spectrum of the nightside of Venus based on gaseous absorption line databases and Mie scattering calculations for cloud particles as sulfuric acid droplets. The model is described by Tsang et al. [2008], with surface emissivity 0.8, surface temperature 740 K, a continuum absorption coefficient of 3.0 109 cm1amagat2, cloud model from Pollack et al. [1993], and no subcloud haze. [24] In windows shortward of 1.74 mm, cloud particles scatter almost conservatively. Single scattering albedos differ from unity less than 103 in magnitude [Grinspoon et al., 1993, Figure 1d]. Subsequently, cloud optical properties are approximately the same at the wavelengths of these windows. Meadows and Crisp [1996] exploit this to diminish cloud contrast in imaging observations via the ratio of surface windows (1.02, 1.10, and 1.18 mm) to the atmospheric window at 1.31 mm. At 1.31 mm the atmosphere becomes strongly absorbing between clouds and surface. Atmospheric brightness temperature can assumed to be horizontally constant below the clouds wherefore observable contrast of flux at top of atmosphere gives local transmittance of flux through the clouds. To interpret thus ‘‘declouded’’ images, Meadows and Crisp [1996] construct a synthetic image based on their model of the transfer of surface thermal emission. Temperatures were calculated from Pioneer Venus altimetry and a uniform emissivity of basalt (0.85) was chosen. The ratio of synthetic to declouded observed image showed less than 10% variation which is presented as upper limit of surface emissivity variation at 1.18 mm. There is still a slight negative correlation with the cloud contrast to be seen in the declouded image. Additionally, for the effect of topography or surface temperature to vanish in emissivity, a lapse rate with 7.5 K km1 has to be chosen, which is different from the 8.06 K km1 given by the VIRA model [Seiff et al., 1985]. [25] As Moroz [2002] and Hashimoto and Sugita [2003] point out, this approach employing simple ratios to account for cloud and topography contrast neglects, that emerging flux at a surface window is neither proportional to surface emissivity nor to the cloud transmittance determined from 1.31 mm because of the different lower boundary conditions. While at 1.31 mm the deep atmosphere below the clouds reflects no downwelling radiation, at 1.02 mm atmospheric absorption is small enough for multiple reflections between atmosphere and surface to be relevant for emerging flux. Because cloud reflectivity and surface albedo are negatively correlated to cloud transmittance and surface emissivity, respectively, the contrast of emerging flux that contains significant contribution from multiple reflections is diminished compared to the direct proportionality implicitly assumed by Meadows and Crisp [1996]. Hashimoto and Sugita [2003] conclude that the slight negative correlation with cloud contrast in the emissivity
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presented by Meadows and Crisp [1996] is likely due to this effect and also that actual variation of emissivity might be higher than reported by Meadows and Crisp [1996]. Hashimoto and Sugita [2003] also employ a two-stream approximation to quantify the effect of multiple reflections and to study the observability of the surface emissivity from orbit. This approximation is here adapted to parameterize the effect of the atmosphere on the emerging thermal emission at 1.02 mm. 2.3. Atmosphere Parametrization [26] Moroz [2002] and Hashimoto and Sugita [2003] give an approximation for the relation between surface thermal and emerging flux Fl at the top of a emission flux Ftherm l highly reflective atmosphere Fl ¼
t F therm 1 al r l
ð1Þ
where t is transmission of the atmosphere, r is reflectivity of atmosphere and al is lower boundary albedo. Equation (1) is exact for plane parallel conservative atmospheres which is approximately true for atmospheric layers composing the clouds at the window wavelengths and it is t = 1-r. Hashimoto and Sugita [2003] assume that cloud reflectivity r is the same for both surface window and the atmospheric window at 1.31 mm. therm is constant because the atmosphere becomes [27] F1.31mm opaque owing to absorption and emits in local thermal equilibrium with the constant atmospheric temperature. Brightness of the thermal emission from radiative transfer modeling [Tsang et al., 2008] gives a rough approximation for the height of the transition from a conservative to absorbing atmosphere and thus for the source region of radiation (Figure 2), for a better estimate see the contribution functions by Tsang et al. [2008]. The isoline of t r = 0.1 in Figure 2 illustrates that Rayleigh scattering is not very important in the source region of the 1.31 mm window at approximately 20 km height. Little downwelling radiation is scattered upward in the source region; this allows us to assume a1.31mm = 0 [Hashimoto and Sugita, 2003]. Equation (1) then becomes r ¼1
F1:31mm therm F1:31mm
ð2Þ
and allows us to calculate to calculate the atmospheric reflectivity r above the source region of the 1.31 mm window. [28] Cloud reflectivity r contains all required information on the variable cloud transmittance and can be used to remove cloud contrast at 1.02 mm with equation (1). Cloud reflectivity r does not account for scattering below the source region of the 1.31 mm window, however there is significant scattering at 1.02 mm below 35 km due to Rayleigh scattering (Figure 2) and possibly due to aerosols [Moroz, 2002; Grieger et al., 2004]. Additionally, there is some small but poorly constrained amount of gaseous absorption due to the high pressure. Therefore, the paramdo not directly represent the surface in eters al and Ftherm l the frame of equations (1) and (2), but rather a composite of surface and lowest atmosphere that cannot be disentangled
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Figure 7 shows that a constant al is consistent with the data and a certain fit of Tl(Z). [30] Equation (3) relates the observable flux Fl to three parameters, upper atmosphere reflectivity r, lowest atmosphere and surface reflectivity al, and temperature parameter Tl(Z), none of which is directly indicative of surface composition. Since surface composition variability is the aim of this study, equation (3) has to be interpreted in light of the implicit assumption of no surface emissivity variability. Any local deviation of VIRTIS data from the solution of equation (3) with empirically found values for the three unknown parameters is related either to local variations in the properties of the atmosphere or, since there is yet no evidence for such variations, to surface properties. To quantify this deviation, VIRTIS data are introduced as Fl and thermal flux anomaly Al is defined as deviation from unity of the ratio of both sides of equation (3) Al ¼ 1
Figure 3. Sketch illustrating the atmosphere parametrization. Layer 1 is the lowest part of the atmosphere, from the surface to 35 km altitude. Its average behavior, including the surface, is described empirically. Source of thermal radiation is assumed to be a blackbody emitting flux p Bl[Tl(Z)]; scattering in the lowest atmosphere is represented by al. Layer 2 is the atmosphere from 35 km to the top of the atmosphere, which is approximately conservative and grey. This allows us to account for the cloud contrast at 1.02 mm using the VIRTIS images at 1.31 mm to determine cloud reflectivity r. Deviation of data from top of atmosphere flux derived using this model may indicate surface temperature or emissivity anomalies. easily. Instead, global average behavior of these two parameters is determined empirically from VIRTIS data in section 4.4. [29] For a better understanding of the parameters al and , consider an additional atmospheric layer (Figure 3, Ftherm l layer 1) below the part of the atmosphere that is spectrally grey and conservative (Figure 3, layer 2). There is no evidence for horizontal variation in optical properties other than surface topography Z in this layer. Surface topography affects extinction optical depth but also determines average surface temperature. For sake of simplicity, the surface is represented by a blackbody with hemispherically integrated thermal emission pBl[Tl(Z)], with a temperature term Tl(Z) that encompasses all brightness variations due to topography. Topography-independent scattering is represented by al. When conservative scattering is assumed, this allows us to rewrite equation (1): Fl ¼
ð1 rÞ ð1 al ÞpBl ½Tl ðZ Þ 1 al r
ð3Þ
Tl(Z) is an approximation for surface temperature. This approximation deviates from real surface temperature owing to absorption, multiple reflections within the lowest atmosphere, and the dependance of optical depth on topography Z. In reality, al will depend on surface topography too, but
Fl ð1 al rÞ ð1 rÞð1 al ÞpBl ½Tl ðZ Þ
ð4Þ
Surface emissivity is positively correlated with Al, but surface temperature variation not expressed by Tl(Z) also affects it.
3. Surface Observations by VIRTIS on Venus Express [31] The VIRTIS instrument on Venus Express is a flight spare of the instrument of the same name of the Rosetta mission [Drossart et al., 2007]. The versatility of the instrument originally designed to observe coma and nucleus of a comet at moderate spectral and spatial resolution [Coradini et al., 1998] allows imaging of all NIR windows longward of 1 mm wavelength from Venus orbit. The infrared mapping subsystem VIRTIS-M IR that acquired the data for this study diffracts a light bundle selected by a slit with on a detector array. Data sets acquired at a given time (frames) therefore have one spatial and one spectral dimension, here referred to as samples and bands with index b. The detector array has 256 samples and 431 bands, in total giving a field of view of 64 mrad, corresponding to approximately a third of the diameter of Venus at apocenter, and a spectral range from approximately 1 mm to 5 mm. A scanning mirror with 256 position allows the instrument to construct hyperspectral image cubes by appending frames acquired consecutively at different viewing angles, thus adding a third dimension of lines. Neither mission nor instrument are specifically designed to observe the thermal emission of the surface but the manifold objectives of the mission requires different viewing geometries and integration times [Svedhem et al., 2007]. The signals of the surface temperature and moreover of surface emissivity are small compared to the reflected sunlight on the dayside and require several seconds of exposure on the nightside to gain sufficient signal-to-noise ratio (SNR). Additionally a correction for variable cloud opacity such as attempted here is an additional source of noise that is especially critical since radiance and SNR at 1.31 mm is even lower. The error in cloud reflectivity derived from 1.31 mm is a major contribution to overall uncertainty in retrieval of surface properties [Hashimoto and Sugita, 2003]. There is an atmospheric
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Figure 4. VIRTIS data spectrum from cube VI0112_01, line 0, sample 0. Dotted line is stray light spectrum Sb scaled to fit radiance I½36;39 . Stray light is removed by subtraction of individually scaled stray light spectra from all data spectra, similar to the approach by Meadows and Crisp [1996]. window at 1.27 mm with comparable properties and better SNR. Unfortunately it is coincident with oxygen airglow in the upper atmosphere [Crisp et al., 1996] and unsuitable for a straightforward determination of cloud transmittance. The long duration of the mission and repeated observations of the same area of surface gives the unprecedent opportunity to combine observations over a long period of time to enhance the time invariant signal of the surface and to offset the poor SNR. [32] The 24 h eccentrical polar orbit with pericenter at approximately 80°N exclusively allows for imaging of the southern hemisphere. Science planning of the mission is described by Titov et al. [2006]. Observation schemes are designated as science cases. All relevant cases are briefly summarized here: The northern hemisphere is only accessible in case 1 observations that are composed of isolated lines of spectra perpendicular to the spacecrafts track. Case 3 observations are apocenter mosaics of either the entire visible southern hemisphere or alternatively of either dayside or nightside. Owing to the eccentrical polar orbit these observations show the equatorial regions only at very high emission angles and accordingly low spatial resolution. Case 2 observations are done on the ascending or descending branch of the orbit and therefore able to observe regions up to 30°S with fairly low emergence. During most observations, especially case 3, nadir is at low latitudes and on average observed latitude is correlated to emission angle. At low latitudes the terminator is in or close to the field of view and observations are affected by sunlight scattered in the upper atmosphere and instrumental stray light originating from the bright side of Venus. In addition to the viewing geometry affecting the measured radiance, there is a shift from observation to observation in band positions due to varying thermal strain in the instrument.
4. Data Processing
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1.0255 mm are processed. A uniform 7.5 nm shift is applied to the wavelengths table in the VIRTIS data cubes to better reconcile data with synthetic spectra. Parameters empirically derived during processing either have the subscript b, giving band number as integer starting at zero, or l giving band position in micrometer. The first pertains to all images of that band regardless of individual band positions while l denotes that the parameter in question is seen as a function of individual band position wavelength of individual observations. [34] Several different VIRTIS bands are situated inside the window regions. The band with the best ratio of surface to atmospheric influence is b = 0, on average at l = 1.021 mm. The band closest to the center of the 1.31 mm is at b = 30 on average at l = 1.307 mm. Analysis of VIRTIS data is performed only for band 0 for the surface signal and band 30 for the cloud signal, although the data processing mechanism may as well be extended to other windows and bands farther from the center (e.g., Figure 4). It is possible to use the average of several or all bands showing a signal from the surface with this approach. This would increase SNR for individual observations and help in the detection of transient temperature variations, e.g., thermal signature of active volcanism. The contrast related to surface emissivity, however, is decreased in the averaged image as the transparency of the atmosphere is greatest in the small wavelength range used here. [35] Only observations with 3.3 s exposure duration or greater have a sufficient SNR ratio. In total 1161 case 2 and 3 observations have been selected for processing. Observations span a period from May 2006 to December 2007. Some observations contain frames, where all spectra of one line are offset relative to the neighboring frames. These frames are not processed and appear as horizontal lines of missing data in VIRTIS images (Figure 9). 4.1. Scattered Sunlight [36] VIRTIS image cubes on the nightside are affected to some extend by sunlight. While sunlight scattered by the upper atmosphere extends to at least 95° of incidence, the bright illuminated crescent of Venus close to the field of view of VIRTIS also causes stray light within the instrument. The same is true for direct sunlight, which is however, due to the geometry of the orbit, less relevant for the case 2 and 3 observations than for case 1 [Arnold et al., 2008]. Meadows and Crisp [1996] use a spectrum extracted from the sunlight crescent as template for the removal of stray light from the 1 mm region. The spectrum is scaled to fit the radiance at the 1.4 mm CO2 absorption band for each of their spectra and then subtracted. [37] In a similar approach here the spectral average of the VIRTIS bands 36 to 39 is used as parameter I½36;39 for the spatial distribution of stray light. To determine the spectrum of stray light, a linear regression is performed over all VIRTIS observations for the radiances of VIRTIS bands 0 to 35 against this parameter assuming a direct proportionality between the stray light seen at I½36;39 and the radiance of other bands Ib. The linear regression retrieves the best fitting parameters a to the equation
[ 33 ] To minimize the influence of the wavelength shift due to instrument temperature, only observations with the first band situated between 1.0175 mm and 7 of 21
Ib ¼ Vb þ Sb I½36;39
ð5Þ
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4.2. Limb Darkening [39] Limb darkening due to scattering in the upper clouds affects all VIRTIS images, most severely in off-nadir case 2 and in case 3 southern hemisphere mosaics. The limb darkening of synthetic spectra agrees well with VIRTIS radiances (Figure 5) and shows that the angular distribution of emitted radiance I0b does not significantly vary between the 1.02 mm and the 1.31 mm window. The angular distribution of radiance with respect to cosine of emission angle x can be described as product of nadir radiance I0b(1) and limb darkening function P(x). Ib0 ðx Þ ¼ Ib0 ð1ÞPðxÞ
ð7Þ
Required for the two stream approximation employed here is hemispherically integrated flux [Goody and Yung, 1989] Fb
¼ 2p
R1
0 0 Ib ðx Þxdx
¼ 2pIb0 ð1Þ
R1 0
ð8Þ
PðxÞxdx
To account for viewing angles, the radiance I0b(x) measured at each pixel is divided by the limb darkening function P(x) determined by interpolation of synthetic spectra at different emission angles scaled to unity at nadir (x = 1), see solid line in Figure 5. The angular integration is performed by approximating P(x) = 0.31 + 0.69x which in conclusion yields Figure 5. Distribution of surface window radiance I0, atmospheric window radiance I30, and corresponding cosines of emission angles x; only data points with negligible stray light I½36;39 0.005 W m2sr1 mm1 are considered. Frequency distribution is scaled to yield one at maximum. Solid line is created from synthetic spectra of the nightside at different emission angles, and dashed lines are function P(x) = 0.31 + 0.69x scaled to fit synthetic spectra at nadir. This separates the flux Ib spectra into two terms, first thermal emission of Venus Vb which is not correlated to flux at bands 36 to 39 but represents average thermal emission, and second stray light spectrum Sb as the coefficients of proportionality of Ib to I[36,39]. The linear relation extends to a value of approximately I[36,39] = 0.01 W m2mm1sr1 and only points in this range are used for the regression and further processing. [38] Comparison of VIRTIS spectra with the stray light spectrum Sb scaled to fit radiance at bands 36 to 39 demonstrates that there is a good fit at absorption bands while at window wavelengths, stray light is slightly increased above the lower envelope of the VIRTIS data spectrum, see Figure 4. Individually scaled stray light spectra are subtracted from each data spectrum to isolate the local flux of thermal emission: Ib0 ¼ Ib Sb I½36;39
ð6Þ
The values relevant for data processed here are S0 = 3.30 and S30 = 1.32.
Fb ¼ p0:77
Ib0 ðxÞ Pðx Þ
ð9Þ
4.3. Projection and Smoothing of Magellan Altimetry [ 40 ] The Magellan global topography data record (GTDR) from the planetary data system (PDS) gives topography of the surface of Venus [Ford and Pettengill, 1992]. For the VIRTIS observations, the data are projected on the VIRTIS field of view and incorporated in the ancillary geometry data files distributed together with VIRTIS hyperspectral cubes. Since the thermal emission of the lower atmosphere and surface is scattered in the clouds, the radiance observed by VIRTIS originates in the cloud layer illuminated by the surface and atmosphere below. This has two consequences: First, the altitude relevant to a pixel is not given by the intercept of the line of sight (LOS) with the surface, but by the altitude directly below the intercept of the LOS with the cloud layer. Second, not only the point directly below LOS cloud intercept but the neighboring area contributes to illumination of the clouds at this point. This results in a smearing of the image of the surface. Simulations by Hashimoto and Imamura [2001] describe this effect as photons from a point source on the surface scattered into a gaussian distribution with full wide half maximum (FWHM) of 90 km above cloud level. Moroz [2002] approximates a spatial resolution of 100 to 200 km. [41] To simulate this effect, the Magellan altimetry in the VIRTIS geometry is projected according to VIRTIS image viewing geometry to a height of 65 km corresponding to the upper main cloud deck. The spatial resolution of VIRTIS images varies on the southern hemisphere between approx-
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0.04%. In some regions, especially rifts, the difference is higher but does not exceed 2%. No strong dependance of the difference to surface altitude has been found, the error in the determination of global brightness temperature is assumed to be small. The width of the smoothing weighting function is likely a more critical parameter but no study of its influence has been made yet.
Figure 6. VIRTIS fluxes F30 binned according to band position l with 1s uncertainty. Solid line is synthetic flux spectrum from plot 1 reduced in spectral resolution by convolution with a gaussian with full wide half maximum (FWHM) of 18 nm and scaled with a factor of 1.36. Assuming cloud transmittance of 0.18 [Hashimoto and Sugita, 2003], this spectrum is used for approximation of spectral dependance of T30. imately 10 and 60 km. The spatial resolution of the projected altimetry has to be aligned with the spatial resolution of the thermal emission flux penetrating the atmosphere. To achieve this task, a moving weighted average is applied. The weighting function is a gaussian with respect to distance from pixel center on the surface and has a FWHM of 100 km. There are some complications to the problem of connecting thermal emission flux and altimetry. [42] First, flux is calculated by angular integration of the blackbody radiation. The hemispherical integration through the factor of p is only exact for plane surfaces, here no attempt has been made to account for curvature of the topography. Second and more important, thermal emission radiance is not linear to surface altitude. Because thermal emission varies strongly with temperature, low altimetry areas contribute unproportionately more to the brightness of a region than higher areas in the same region. On the other side atmospheric transmittance at highland regions is higher. Averaging over altimetry instead of thermal flux leads to an error depending on distribution of altimetry values within the smoothing radius; it is most severe in regions with high slopes and altitude differences. [43] To minimize this error without detailed simulations of radiative transfer, the smoothing algorithm is used in two different ways. For the determination of global atmosphere and surface parameters (section 4.4), the moving average is applied directly to altimetry. Thus, smoothed Magellan altimetry is denoted by Z in units of meters relative to mean planetary radius (MPR) of 6051.84 km [Ford and Pettengill, 1992]. For the calculation of local flux anomaly (section 4.5), the thermal emission flux is calculated on the basis of global brightness temperature to altimetry relation 8 and the original Magellan altimetry projected on the VIRTIS field of view. The moving average is then applied to the thermal flux. [44] The difference between flux calculated from the smoothed altimetry, and the smoothed flux is on average
4.4. Atmospheric Parameter Determination [45] Equation (3) relates emerging flux at surface window to three parameters: r, al, and Tl(Z). Typical cloud reflectance of r = 0.82 based on a nominal cloud model is given by Hashimoto and Imamura [2001]. Adopting this value in = 0.77pBl[Tl] leads to equation (2) and assuming of Ftherm l a temperature of the atmosphere below the clouds Tl. Shape of the synthetic spectrum (Figure 1) is used to fit spectral dependance of Tl to average VIRTIS flux at different wavelengths (Figure 6). The resulting range of T30 from 528 K to 531 K approximately corresponds to VIRA atmospheric temperature at 26 km height for all wavelengths covered by band 30. Flux F30 measured with VIRTIS band 30 and estimated T30 allow to calculate local cloud reflectivity r with equation (2). [46] The two remaining parameters are estimated from VIRTIS and Magellan altimetry data. While atmospheric temperature at the mean planetary radius of 735 K and surface emissivity e0 > 0.8 are relatively well constrained by in situ measurements, adopting these values as e0 = 1 a0 and T0(0 km) = 735 K does not work well because lowest atmosphere scattering and absorption is not accounted for. Albedo a0 is expected to be greater than 0.15, which Meadows and Crisp [1996] used to represent the basaltic plains of the surface. Temperature T0 with respect to equation (3) will be less than real surface temperature since the lowest atmosphere is not fully transparent. [47] To estimate a0, the distribution data points in the F0, F30, and Z space is determined using all VIRTIS observations. Average value of F0 in each F30 interval is given with 0.5s error bars for different topography intervals in Figure 7. The position on the F30 axis translates with equation (2) into cloud reflectivity r. Figure 7 demonstrates that it is possible to reproduce the relation of observed fluxes F0 and F30 within 0.5s with a lowest atmosphere and surface albedo a0 = 0.22 independent from topography and temperature T0(Z) as monotonic function of altitude. Variation of a0 with band position l has not been investigated. Preliminary analysis of the windows at 1.10 and 1.18 mm not presented here yields similar values of lowest atmosphere albedos, which leads to the conclusion that this parameter does not vary strongly with wavelength. [48] Using this constant, a0, equation (3) is solved for T0 for each data point. Next, the distribution of T0 values in the topography Z and band position l space is determined from all processed data points. A two dimensional polynomial of degree 2 with coefficients knm and parameters Z and l is fitted to the averages of T0 to obtain an analytical approximation of Tl(Z):
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the following conditions are met: (1) emission angle is less than 80°, (2) incidence angle is greater than 95°, and (3) sunlight parameter I½36;39 is less than 0.01 W m2mm1sr1. [51] Two different quantities are retrieved, either surface temperature T0 or the thermal flux anomaly A0 which is thermal emission corrected for the global relation of surface temperature Tl(Z) to altimetry and band position. Figure 9 illustrates the process of retrieving surface and lowest atmosphere brightness temperature. Both results are map projected for each observation using Lambert’s azimuthal equal area projection centered on the south pole. The projected individual images are then combined by using the median of all different observations of the same place. The spatial distribution of processed images is shown in Figure 10.
5. Results Figure 7. Average of surface window flux for different altitudes as a function of cloud reflectivity r. Flux uncertainty is given as 0.5s. Size of topography bins is 500 m. Curves are calculated with equation (3) with a0 = 0.22 for all heights and T0(1 km) = 712 K, T0(0 km) = 708 K, T0(1 km) = 703 K, and T0(2 km) = 698 K. Figure 7 shows that equation (3) describes the relation between surface and atmosphere window flux well and thus can be used for removal of cloud contrast. The i j matrix of coefficients is 0 B B kij ¼ B B @
63:6 103
12:3 103
60:3 103
13:0
25:6
12:5
1:83 103
3:62 103
1:79 103
1 C C C ð11Þ C A
Units of kij are mjmmi K where i is index of column and j is index of row starting with zero. [49] The Z-l isolines determined directly from the data and from the fit of equations (10) and (11) are illustrated in Figure 8 for seven temperatures. The lapse rate of 4 to 5 K km1 is not consistent with the 8.06 K km1 in the VIRA model but this is not surprising since scattering, absorption and emission in the atmosphere lowers the lapse rate of observed flux. Furthermore scattering and absorption optical path from surface to top of atmosphere is determined by topography. Usually this is negligible, as topography is only a fraction of total optical path lengths, but extinction is here strongly concentrated toward the lowest part of the atmosphere owing to Rayleigh scattering being proportional to density and the continuum absorption being proportional to square of density [Pollack et al., 1993]. 4.5. Image Processing [50] After the parameters necessary for atmospheric correction are either assumed or determined from the statistical behavior of VIRTIS data points, the next step is the application of equation (3) on each observation to analyze the local thermal flux. Data pixels are processed only when
[52] Surface temperature parameter T0 averaged over time is highly correlated with altimetry in Figure 11. Signal to noise is better in the hemisphere west of 0° longitude as is expected owing to greater coverage (Figure 10). SNR generally decreases toward the limb, where less images as well as less spectra per unit area have been acquired owing to the generally oblique viewing angles. The polar region is usually seen at low emission angles but stray sunlight increases the error although spectra surpassing a certain threshold of stray light are discarded during processing. [53] The map of flux anomaly Figure 12 presents in principle the same data as Figure 11 but corrected for the global temperature to topography relation (Figure 8). There are however some correlations with topography remaining when comparing the altimetry contours in Figure 11 with flux anomaly in Figure 12. There is lower flux in the sector between 120° and 110° longitude starting at about 75° to 60°S latitude and extending at least to the equator. This coincides with an increased uncertainty in the Magellan altimetry investigated by Rappaport et al. [1999]. Several rifts, e.g., Artemis Chasma in V-48, show contrast in the flux anomaly map. This might be related to the difficulties in smoothing of the radar altimetry (section 4.3). The error
Figure 8. Relation of lower atmosphere surface brightness temperature T0 (for seven different values of T0) to topography Z and band position l. Solid lines are isolines of average data temperature. Dashed lines are polynomial fit to data.
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Figure 9. VIRTIS image VI0373_01 (second image acquired in orbit 373) during several processing steps: (a) band 0 at 1.02405 mm; (b) spectral median of bands 36 to 39 showing distributions of stray light; (c) band 0 corrected for stray light and limb darkening; (d) band 30 corrected for stray light and limb darkening; (e) brightness temperature of lower atmosphere and surface derived from band 0 and band 30; and (f) surface temperature map constructed from Magellan altimetry and Tl(Z). Horizontal dark lines are data points not processed owing to an offset of the whole frame. due to the smoothing is expected to be greatest in regions with steep slopes and large altitude differences. [54] The equatorial regions generally show a greater residual correlation of flux and topography. This might be indicative of a latitudinal component in the flux to topography relation, which might be due to latitudinal temperature or atmospheric composition variation. Most data points used for determination of the global flux to topography relation are in the midlatitude to low-latitude regions between 25° and 75°S where influence of topography seems to be well accounted for (compare Figure 10 and Figure 12). Spectrometer temperature and thus band position l varies
with orbital position. The equatorial regions have a bias to be observed at wavelengths displaced from the optimum. Determination of Tl(Z) is therefore possibly less reliable for the equatorial regions. The lower latitudes are, with some exceptions that are described in sections 5.1 – 5.3, very well represented by the global flux to topography relations as the flux anomaly is very near to zero. 5.1. In Situ Sites [55] There are four in situ measurements of surface rock composition on the southern hemisphere, see Figure 12. The Venera 8 landing site in quadrangle V-43 is in a region with
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Figure 10. Spatial distribution of processed images in Lamberts azimuthal equal area projection centered on the south pole and extended to the equator; labels denote Venus mapping quadrangles. average to low thermal flux. Venera 13 in quadrangle V-42 is in the region with low flux adjacent to Doliya tessera, Venera 14 is on the gradual transition from this same diffuse low flux area to the diffuse high flux area over Navka planitia. Vega 2 landed on Rusalka planitia in quadrangle V37 which shows an unreliably determined flux. All landing sites are in the equatorial regions and not very well covered by VIRTIS observations. Possible interpretations of rock types are summarized on the basis of the work of Kargel et al. [1993]. Venera 8 measured abundances of natural radioactive isotopes of K, U, and Th, which where consistent with several terrestrial analogues, including rhyolite (volcanic rock of felsic composition), monzonite (intermediary composition) and leucitite, an on Earth rare alkaline mafic volcanic rock in which felsic minerals are mostly replaced by the feldspathoid mineral leucite. Leucitite is also proposed for the Venera 13 landing site, where elemental abundances were more directly measured by X-ray fluorescence spectroscopy. The Venera 14 and Vega 2 samples also analyzed with XRF spectroscopy are interpreted to have basaltic composition similar to mid-ocean ridge basalts. Though it might be tempting to attribute areas of low flux anomaly to the exceptional samples of Venera 8 and Venera 13 situated in the vicinity, very little difference is seen in NIR flux anomaly between the Venera 8, Venera 13, and Venera 14 sites. 5.2. Relation to Geological Settings [56] The thermal flux anomaly shows some correlation with geomorphological terrain types and also individual features. Over large areas of tessera terrain the flux adjusted for topography is generally lower than in neighboring areas. The tessera terrain outlines in Figure 12 are taken from the
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electronic global geological map accompanying the work of Tanaka et al. [1997]. Large tessera terrains with relative good SNR are Phoebe regio and Doliya tessera in V-41, Alpha regio in (among others) V-32, Lhamo and Cocomama tessera in V-56. Tessera boundaries are however not well defined in NIR flux, negative flux anomalies are interspersed with average flux and occasional positive anomalies. For instance, at Doliya tessera no relation of its boundaries to flux is obvious, but instead the tessera area is situated directly on the boundary between larger areas of positive and negative flux anomalies. [57] The negative flux anomaly over tessera regions is consistent with the relative low surface emissivity at highlands retrieved by Hashimoto et al. [2008] from the Galileo NIMS flyby data. Tessera are more frequent in highlands [Ivanov and Head, 1996]. Their approach does not correct for topography empirically but uses a flux to topography relation derived from the adiabatic lapse rate and radiative transfer modeling and is therefore not directly comparable to this result. [58] Tessera are not extensively covered by VIRTIS observations; there is very little data yet of the large tessera highlands in Aphrodite terra. Most VIRTIS spectra were acquired on midlatitudes to low latitudes <25°S between 180° and 360°E longitude, dominated by tectonically modified lowland plains. Highlands in this area are Imdr, Ishkus, Thetis, and Dione regios where some, but not all, large volcanic edifices show a flux exceeding average flux at comparable altitude by 5%, listed in Table 1. Two coronae are also listed. The area of these flux anomalies is significantly smaller than the topographic rise and concentrated on the flank. [59] A similar but larger anomaly is found at the southern flank of the Quetzalpetlatl corona rise in Lada terra, 0°E 70°S. The southern boundary of this anomaly is not clearly defined as it borders on the polar area with small SNR ratio. The fan shape nevertheless correlates with the large lava flows Juturna and Cavilaca fluctus, extending from the rim of Boala corona, nested inside Quetzalpetlatl corona. For most recent radar imaging see the work by Kratter et al. [2007]. The contours of digitate plains units from the electronic material accompanying Tanaka et al. [1997] are outlined in Figure 13. Ivanov and Head [2006] characterized the unit containing the positive flux anomaly south of Quetzalpetlatl as tectonically undeformed lobate plains. Lavinia planitia in V-55 shows an overall increased flux compared to other planitia regions at altitudes below mean planetary radius and especially prominent at the eastern basin rim. This basin rim is characterized by several lobate lava flows and plains (Figure 13), emanating from coronae, volcanoes or the Lada rift [Magee and Head, 1995]. These units are designed stratigraphically young, undeformed lobate plains in the map by Ivanov and Head [2001]. The spatial correlation is, however, less significant than at the lava streams south of Quetzalpetlatl and the anomaly does not exceed 10%. For a more detailed analysis of the NIR flux anomaly at Quetzalpetlatl see Helbert et al. [2008]. 5.3. High-Altitude Radiothermal Anomaly [60] Radiothermal emissivity as measure by the Magellan mission [Pettengill et al., 1992] distinctly drops off from a value around 0.85 to a value of 0.4 above surface elevations
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Figure 11. Surface lower atmosphere brightness temperature T0 mosaics in Lamberts azimuthal equal area projection centered on the south pole. Magellan altimetry is plotted with 1 km contours. Black hachures denote missing altimetry data. Temperature is well correlated with altimetry, although influence of wavelength shift is not equal in all regions. of 4 to 5 km [Wood, 1997]. This height is rarely reached in the area mapped by VIRTIS with good SNR. Areas with radiothermal emissivity less than 0.7 are marked on the flux anomaly map in Figure 12 by dark hachures from up left to down right. A single region with a radiothermal emissivity anomaly, quadrangle V-41, Yunya-mana mons (285°E, 18°S) is mapped with reasonable SNR and shows no significant flux anomaly in NIR. However, the extent of this area is close to spatial resolution of the NIR data and observed flux might be influenced by the neighboring tessera terrain. The Venus Monitoring Camera (VMC) also on Venus Express did not observe resolvable emissivity variations in this region (Basilevsky et al., submitted manuscript, 2008). [61] Repeated coverage of large regions showing a radiothermal emissivity, e.g., Thetis regio in V-36, is required to analyze the radiothermal anomaly in NIR. The approach used here is however not very well suited for this investi-
gation. The influence of any present altitude-dependent variation of emissivity will be to some extent removed during the processing owing to the empirical determination of Tl(Z). A nonempirical approach that correctly accounts for absorption and scattering independent from VIRTIS data in the lowest atmosphere, as in the work of Hashimoto et al. [2008] or Arnold et al. [2008], is required. Data at higher latitudes and altitudes, e.g., Maxwell Montes are also desirable, but will not be available in observations other than case 1 northern hemisphere sparse imaging.
6. Discussion [62] Interpretation of the thermal flux anomaly is highly ambiguous. First, a flux anomaly due to surface emissivity cannot be distinguished from surface temperature effects with this approach. The long period of one and one half year over which the data are averaged and the properties of the
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Figure 12. Map showing flux anomaly A0 in relation to in situ measurements and tessera terrain. Landing site coordinates of Venera and Vega probes are from Basilevsky and Head [2003]. Tessera terrain outlines are from the electronic material accompanying the article by Tanaka et al. [1997]. Relation to in situ measurements is inconclusive. Tessera show a negative flux anomaly. Flux over tessera terrain is less than in plains and volcanic rises of the same Magellan altimetry values. Only areas with a coverage of three or more VIRTIS images are mapped here. lower atmosphere of Venus (see section 2.1) preclude, to first approximation, local surface temperature variation due to atmospheric or insolation effects. Endogeneous heat, e.g., active volcanic vents or even lava flows, could also be a source of thermal NIR flux and the coincidence of positive flux anomalies and relatively young volcanic features presents this as an attractive explanation. However, the recent rate of extrusive volcanism is comparable to intraplate volcanism on Earth [see Grimm and Hess, 1997, and references therein]. Cooling heat flux of liquid lava on the surface is slightly lower than on Earth [Snyder, 2002], but cooling time scale of the lava is estimated to be in the order of 1 day; that is, an eruption is not detectable for more than 1 day after its end [Hashimoto and Imamura, 2001]. Areas of increased flux are up to the order of hundreds of
Table 1. Volcanic Edifices Correlated With Positive Flux Anomalya Diameter (km) Name Mertseger mons Shiwanokia corona Shulamite corona Idunn mons Hathor mons Ininni mons
Coordinates 270°E, 278°E, 284°E, 215°E, 324°E, 328°E,
37°S 42°S 39°S 46°S 38°S 34°S
Anomaly
Structure
125 200 200 150 250 125
450 500 275 250 333 339
a Names, coordinates, and diameters of volcanic edifices are from the Web page of the USGS Astrogeology nomenclature database (available at http://planetarynames.wr.usgs.gov/). Two coronae are included in this compilation; their flux anomaly is located on the flank of the topographic rise but outside the annulus.
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Figure 13. Map showing flux anomaly A0 in relation to plains with a flow-like morphology and to the radiothermal emissivity anomaly in high regions. Digitate plains outlines from the material accompanying Tanaka et al. [1997]. Correlation of radiothermal end NIR emissivity is inconclusive owing to the insufficient coverage and the nature of the topographic correction. Most prominent coincidence of positive flux anomaly and digitate plains is on the southern flank of Quetzalpetlatl corona, 0°E and 70°S. This plains unit is composed of the large lava flow fields fan shaped extending from Boala corona on top of the topographic rise [Ivanov and Head, 2006]. thousands of square kilometers, temporal sampling rate is at most one or two images per day, typically much less. The map is a temporal median of up to several tens of images, and therefore filters the effects of short or even singular events with significantly increased temperatures. All individual observations (i.e., the result presented in Figure 9) have been searched for clear evidence of an active volcanic eruption but none has been found yet. Volcanic vent activity can be continuous but is also difficult to reconcile with the extend of the flux anomaly. Owing to these considerations, here flux anomaly is interpreted in terms of surface emissivity. [63] Second, interpretative ambiguity is that variation of surface emissivity known at only one wavelength is difficult to assign to a certain surface material. As only one point in
the surface spectrum is spatially mapped, the emissivity is interpreted as a result of relative ratio of felsic to mafic mineral abundances. Felsic minerals (i.e., feldspars and quartz) generally have a low emissivity at 1 mm while mafic (Mg- and Fe-rich) minerals tend to high emissivities, which would allow to distinguish between felsic granite and mafic basalt [Hashimoto and Sugita, 2003]. The poorly known chemical weathering environment complicates the interpretation of surface emissivity with respect to composition of terrestrial analogues. Minerals affecting emissivity such as pyrite or magnetite might be unstable depending on surface temperature and atmospheric redox state [Fegley et al., 1997; Hashimoto and Abe, 2005]. Wood [1997] relates the high-altitude radiothermal emissivity anomaly to volatile
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transport of magnetite from the hot lowlands to the mountaintops. This would affect NIR emissivity as well but no evidence has been found by Lecacheux et al. [1993], Meadows and Crisp [1996] or Basilevsky et al. (submitted manuscript, 2008). Owing to the empirical nature of this approach, it is very difficult to investigate any present altitude-dependent emissivity trend, though a comparison of VIRTIS data of comparable high altitudes with and without radiothermal anomaly would be very interesting but difficult owing to the orbit of Venus Express. [64] Weathering age might affect surface emissivity independently from altitude. If emissivity constantly decreases with time, e.g., by slow and steady volatile transport of magnetite from the lowlands to the highlands above 4 km height [Wood, 1997], the observed thermal flux anomaly will denote differences in surface age. Increased thermal flux is more frequently found at tectonically undeformed and probably younger areas as opposed to the lower flux over the stratigraphically old and deformed tessera. However, the small thickness of material contributing to NIR surface emissivity dictates that the weathering effects in emissivity take place relatively rapid and then reach an equilibrium state. The large areas with increased emissivity would require that weathering time scale is not small against volcanic resurfacing time scale. From a very general and preliminary view on stratigraphic age, a weathering process that lowers emissivity over time scales of the average age of the surface of about 0.5 Ga is consistent with observed flux anomalies. Weathering should not be disregarded as possible explanation for the flux anomalies. Further hypotheses, as to which kind of surface material regardless of any weathering effects might be responsible for observed flux anomalies, are presented in sections 6.1 and 6.2, and their plausibility is discussed. 6.1. Relative Felsic Composition of Tessera Terrains [65] In the plains, a basaltic composition is supported by the low viscosity inferred from radar images and in situ measurements. No direct morphological evidence for composition and no in situ measurement exists for tessera terrain, which is characterized by complex tectonic deformation patterns. The crater frequency of tessera terrain is not distinctly different from that of plains, but this does not necessarily imply same age since tessera terrain might have undergone tectonic deformation that reset the crater history. At most boundaries tessera terrain is embayed by neighboring plains and therefore stratigraphically older as well as elevated in topography when visible. It more frequently occurs at higher surface altitudes [Ivanov and Head, 1996]. Several plateau shaped highland regions are composed mostly of tessera terrain. These highlands generally show a small geoid to topography ratio which implies shallow depth of compensation due to either density or thickness variations of the crust [Smrekar and Phillips, 1991]. [66] On the basis of these properties, tessera may represent the granitic cratons that Taylor and Campbell [1983] hypothesized to exist on Venus. If tessera terrain were indeed the venusian analogue to archean continental nuclei on Earth the implications would be far-reaching. Under current dry surface conditions and stagnant lid tectonic regime creation of silicic (i.e., granitic) magmas is unlikely since subduction of water rich crust is required [Taylor and
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Campbell, 1983]. D/H ratio measured in the atmosphere of Venus is consistent with evaporation of a primordial ocean of several tens of meters deep and subsequent preferential loss of hydrogen by atmospheric erosion [Donahue and Russell, 1997]. Under cooler and wetter conditions the accordingly different temperature and rheological contrasts in the mantle and crust might have resulted in continuous plate tectonics [Solomatov and Moresi, 1996; Stein et al., 2004]. The geochemical signature of such a hypothetical setting predating the current greenhouse climate era and stagnant lid regime is most likely to be preserved in tessera terrain and would manifest as low emissivity at 1 mm due to the low content of FeO and high SiO2 content [Hashimoto and Sugita, 2003; Hashimoto et al., 2008]. [67] The qualitative result of relative low thermal emission flux above tessera terrain does not dictate the conclusion of granitic surface composition. Absolute surface emissivity is not known and not also very conclusive even if known at only one wavelength. Other materials can have a significantly lower emissivity at 1 mm than basalt. Anorthosite rock is to 90 to 100% composed of the feldspars anorthite and albite, which both have reflectances at 1 mm of about 0.8. Anorthosite maybe even more plausible an explanation for low emissivity regions than granitic composition since its creation has less prerequisites. Anorthositic crust is supposed to differentiate on the top of a fully molten planetary mantle of terrestrial composition [Taylor, 1974]. In this case the situation on Venus would resemble the lunar crustal dichotomy of bright anorthosite highlands and dark basalt mares [Nikolaeva et al., 1992]. [68] The negative flux anomaly observed here is not perfectly aligned with tessera boundaries, see Figure 12. This does not preclude an interpretation of tessera terrain to be mainly composed of less dense felsic material. Basaltic material might have been emplaced by volcanism, accretion of terranes or obduction on the buoyant and thickened crustal block. The basaltic material then would have undergone the same tectonic deformation history characteristic of tessera while the isostatically elevated topography prevented burial by lava floods. On the other hand buried granitic material might undergo remelting leading to felsic volcanism outside of tessera boundaries. [69] However, there are serious concerns to the hypothesis of felsic tessera terrain. Aeolian transport of crater ejecta is known to occur on Venus. Finer grained ejecta and volcanic ash will be transported over larger distances. It is difficult to imaging that over time period of several Ga, that are required for either anorthosite or granites hypotheses, the tesserae had not been covered by a blanket of material closer to average composition. Basilevsky et al. [1997] gives an resurfacing rate of 0.1 to 10 cm per Ma from the fading of radar bright decimeter to decameter ejecta. Though higher winds at higher altitudes might sweep the tessera highlands [Basilevsky et al., 1992] and tectonic deformation may expose the rock below compacted soil, compositional mixing of surface materials will be more relevant for areas significantly older than the relatively young plains and volcanic features. [70] A second and more serious concern is that radar altimetry seems to be less reliable over tessera terrain, with a bias toward lower than actual surface altitudes. Figure 14 shows Magellan SAR mosaic and global topography data
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formed plain slightly northeast of the center of the image, is partly depressed relative to the plains. Emplacement of the plains without covering the depressed terrain seems improbable. Further support for the hypothesis of systematically too low radar altimetry is given by the negative NIR flux anomaly, that is in shape similar to the suspicious depressions, see Figure 14. The eastern, elevated block of tessera shows nonuniform flux anomaly. Isolated depressions circling the southern and western rim of the elevated tessera block have lower altimetry readings than embaying plains. Assuming that these readings are too low, smearing of Magellan altimetry with a 100 km FWHM gaussian weight function during data processing will lead to a diffuse, crescent shaped negative flux anomaly along the southern border of the tessera. Such a flux anomaly is observed. Similar considerations can be made for the other negative flux anomalies in the area. Magellan altimetry of tessera terrain needs to be examined more closely before any definite conclusions on emissivity of tessera can be made. [71] Independent examination with a different approach analyzing VMC images did not reveal NIR contrast in the shape of tessera terrains (Basilevsky et al., submitted manuscript, 2008). Such contrast would be expected if these terrains were fully composed of anorthositic outcrops with an emissivity of 0.5.
Figure 14. Qualitative comparison of different data sets in the Cocomama tessera region, sinusoidal projection from 66° to 60°S and 13° to 30°E. (a) Magellan left-look SAR (image data from NASA/JPL/Magellan). (b) Magellan GTDR altimetry (image data from NASA/JPL/Magellan). (c) VIRTIS thermal flux anomaly. Outlined is the boundary between mostly pristine plains and areas dominated by tectonic deformation. Deformed terrain is mostly tessera, the central and southeastern region with predominant north – south strike is characterized in electronic maps by Tanaka et al. [1997] as ridge belt. Magellan GTDR possibly has a bias to too low values over deformed areas. tesseraSAR-GTDR. record GTDR of Cocomama tessera. Both seem difficult to reconcile with each other. The easternmost part of the tessera terrain is elevated above the surrounding plains. The deformed terrain encircling the smooth, nearly unde-
6.2. Large Scale Ultramafic Volcanism [72] The high thermal flux over young volcanic flows associated with coronae in Lada terra and less clearly over the nearby Lavinia basin might indicate a composition more mafic than the average plains, which are here used to define normal thermal flux. Most evidence, from morphology and in situ measurements, hints to a basaltic composition of plains [Grimm and Hess, 1997]. Basalts are created by partial melting of ultramafic mantle material, the amount of refractory mafic minerals in the liquid phase is determined by temperature and pressure conditions. The cratering record indicates that the most of the plains were created during an episode of increased resurfacing with little magmatic activity afterward [Schaber et al., 1992; Strom et al., 1994; Basilevsky et al., 1997]. Parmentier and Hess [1992] furthermore propose that the upper mantle below the crust is depleted of felsic minerals and FeO through basalt generation during the resurfacing event. The depleted mantle material is assumed compositionally buoyant and not partaking in convection. Head et al. [1994] argue that pressure release melting in mantle diapirs at the base of this layer might produce large volumes of ultramafic, MgO-rich magma that might account for the large lowviscosity lava flows. [73] Terrestrial analogues are komatiite or picrite that have been proposed by Komatsu et al. [1993] as lavas that could thermally erode basaltic surface and thus produce the incised sinuous channels found, e.g., on the young lava flows in or close to Quetzalpetlatl corona. These sinuous rilles are north of and with flow direction away from the thermal flux anomaly seen on the southern flank of Quetzalpetlatl. The mapping of Komatsu et al. [1993] does, however, not extend below 70°S which is unfortunately also the northern rim of the flux anomaly. The lava flows on the rim of Lavinia planitia feature simple and complex channels
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with flow margins that share some characteristics with sinuous rilles but do not necessarily require thermal erosion [Baker et al., 1997]. [74] The Lada-Lavinia regio, which features the most prominent thermal flux anomalies, has been characterized by Magee and Head [1995] as geological setting favorable for the generation of melt. Mantle downwelling below the Lavinia planitia basin is thought to be associated with the extensional Lada rift on the southern and eastern rim of Lavinia [Baer et al., 1994]. The lava flows with anomalous IR flux are in the vicinity of coronae. Coronae are interpreted to be imprints of diapir activity and possibly linked to extensional rifting [Stofan et al., 1997]. Smrekar and Stofan [1997] modeled the range of topographic signatures of coronae as result of an evolutionary process starting from an upwelling mantle plume, its cessation and as latest stage lithospheric delamination. Elkins-Tanton et al. [2007] modeled possible melt compositions in the context of lithosphere delamination and found, among others, ultramafic melts. [75] Constraints on composition further than the assumption of a higher emissivity than that of basalt cannot be made with this simple atmospheric correction model, wherefore it might be idle to speculate on exact surface rock composition at this point. Furthermore the positive flux anomaly over Lavinia planitia is to some extent correlated with altimetry that given the empirical nature of this approach, rouses suspicions of systematical errors due to the assumption of an incorrect temperature. A similar error might pertain to the flux anomaly seen at volcanic edifices with a diameter in the magnitude of the assumed spatial resolution. Assumption of a lower spatial resolution, i.e., smoothing of the radar altimetry with a wider radius will lower the flux anomalies seen at these places. [76] This is not true for the anomalies seen at the flank of topographic rises. The flux anomaly at the southern flank of Quetzalpetlatl is below 70°S, where VIRTIS images generally are affected by stray light. Also in considering the coverage plot (Figure 10) together with the usual viewing geometry of VIRTIS (the slit oriented parallel to the terminator) it becomes obvious that the Lavinia region has a certain bias to be observed at certain samples of the detector array. If the distribution of stray light on the detector array is not homogeneous the removal of stray light might lead to an systematic error that is not removed by averaging because the coverage of the southern hemisphere is not homogeneous. Further collection of data might resolve this issue.
7. Conclusions [77] The atmospheric correction employed here and more importantly, the capability of VIRTIS on Venus Express to observe the same surface area repeatedly, provide consistent indications that some brightness variation in the 1.02 mm window cannot be accounted for by cloud transmittance and surface temperature depending only on altitude. Contrary to previous results in studies by Lecacheux et al. [1993] and Meadows and Crisp [1996], significant contrast, uncorrelated to cloud opacity and surface altitude, remains after atmospheric correction. The above statement owes much, if not everything, to the higher spatial resolution and area
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coverage achieved through the VIRTIS data set. The method employed here does not allow us to retrieve absolute surface emissivity, as it does not accurately address radiative transfer in the lowest atmosphere [Hashimoto and Sugita, 2003]. Furthermore the empirical determination of thermal flux to topography relation may lead to an error depending on distribution of surface emissivity with surface elevation; see Hashimoto et al. [2008] for evidence for such a distribution. [78] Regardless of these intrinsic uncertainties, a relation exists between NIR flux and morphological units from Magellan radar images at medium altitudes. The lower than average flux over tessera terrain is possibly indicative of felsic composition such as granite or anorthosite, which have been suggested to exist on Venus by Taylor and Campbell [1983], Nikolaeva et al. [1992], and Hashimoto and Sugita [2003]. A similar observation and interpretation based on Galileo NIMS data is made by Hashimoto et al. [2008], who find generally lower emissivity in the highlands. A systematical bias toward lower altitude values in the Magellan topography data set over the highly tectonized tessera terrain is also conceivable and consistent with this result. Comparison with other altimetry data sets, e.g., Pioneer Venus, Venera 15/16, and possibly topography derived from Magellan stereo radar imaging might help to resolve this issue. [79] Higher than average flux is found at some volcanic edifices in Lada terra and Themis, Imdr, and Phoebe regions, the latter three classified as active hot spots by Smrekar et al. [1997]. Lavinia planitia is characterized by a anomalously high flux likely caused by systematic errors. The positive flux anomalies located on the flanks of Shulamite, Shiwanokia, and Quetzalpetlatl coronae, are difficult to attribute to any systematical problem. The Quetzalpetlatl anomaly is well correlated with the young lava flows Cavillaca and Juturna flucti. The high flux anomaly at young lava flows is consistent the hypothesis of large volume komatiite or picrite volcanism predicted by Head et al. [1994] as consequence of chemical differentiation of the upper mantle in the course of secondary crust formation. Another possible interpretation would be that lithospheric delamination in association with the Quetzalpetlatl-Boala corona formation [Smrekar and Stofan, 1997] produced the ultramafic lavas [Elkins-Tanton et al., 2007] that lead to the increased NIR flux. [80] Overall probability of these compositional interpretations is however difficult to estimate, as no absolute surface emissivity is retrieved. To retrieve emissivity, an improved model of radiative transfer has to be implemented similarly to the work of Hashimoto et al. [2008] or Arnold et al. [2008]. Crucial input parameters of radiative transfer models, e.g., continuum absorption coefficient, subcloud haze and even surface temperature are poorly constrained. Analysis of VIRTIS data together with Magellan altimetry data might improve knowledge on these parameters [Carlson et al., 1993a]. When a reliable estimate of Venus surface emissivity is achieved, the next step has to be comparison with near infrared emissivity or laboratory samples measured at high temperatures. Little data exist yet on high temperature emissivity of minerals, but more will be hopefully available soon [Maturilli et al., 2007; Helbert and Maturilli, 2008]. This comparison will constrain surface rock
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composition if near infrared emissivity is not dominated by chemical weathering and volatile transport of minerals. In this case some knowledge on the nature of the high-altitude radiothermal emissivity will be gained if data coverage can sufficiently be extended to these regions. [81] Acknowledgments. We thank S. Smrekar for an insightful discussion of the geology of Venus and also for kindly providing many suggestions for the improvement of language. We acknowledge the financial support for the VIRTIS instrument from ASI and CNES. A significant part of this work was made possible by a grant from the Wernher von Braun foundation (Germany). We are grateful to the whole VIRTIS on Venus Express team for providing this excellent data set.
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P. Drossart and S. Erard, LESIA, Observatoire de Paris, 61 Avenue Observatoire, F-75014 Paris, France.
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G. L. Hashimoto, Laboratory for Earth and Planetary Atmospheric Science, Department of Earth and Planetary Sciences, 1-1 Rokkodai-cho, Kobe University, Kobe 657-8501, Japan. J. Helbert and N. Mueller, Institute of Planetary Research, German Aerospace Center, Rutherfordstrasse 2, D-12489 Berlin, Germany. (
[email protected]) G. Piccioni, IASF-INAF, via del fosso del cavaliere 100, Rome I-00133, Italy. C. C. C. Tsang, Atmospheric, Oceanic, and Planetary Physics, Department of Physics, University of Oxford, Clarendon Laboratory, Parks Road, Oxford OX1 3PU, UK.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, E00B10, doi:10.1029/2008JE003087, 2008
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Venus surface data extraction from VIRTIS/Venus Express measurements: Estimation of a quantitative approach Gabriele Arnold,1,2 Rainer Haus,3 David Kappel,2 Pierre Drossart,4 and Giuseppe Piccioni5 Received 24 January 2008; revised 17 April 2008; accepted 15 July 2008; published 7 October 2008.
[1] Nightside emission measurements of the Visible and Infrared Thermal Imaging
Spectrometer (VIRTIS) on the Venus Express spacecraft were used to estimate the potential for surface data extraction. A selection of orbits over the northern hemisphere was performed for footprints that cover different scales of surface elevation variations. A correction method was used to remove stray light from the measured spectra that is due to direct sunlight striking the instrument. A preliminary radiative transfer calculation technique was applied to simulate Venus nightside radiation. The basic features of the measured spectra are well reproduced. Present limitations of the algorithm are discussed. The variability of the emission window radiances with respect to cloud opacity and surface elevation is modeled and discussed in direct comparison with the measurements. It is demonstrated that a multispectral analysis in the surface and deep atmosphere window ranges (1.0–2.3 mm) and the use of radiance ratios are well suited to decloud the data and to extract surface information from the VIRTIS measurements. This method allows a mapping of surface topography and the retrieval of the surface temperature. A preliminary topography retrieval at Beta Regio was performed and compared with Magellan radar data. Differences are possibly due to emissivity variations on the surface. Citation: Arnold, G., R. Haus, D. Kappel, P. Drossart, and G. Piccioni (2008), Venus surface data extraction from VIRTIS/Venus Express measurements: Estimation of a quantitative approach, J. Geophys. Res., 113, E00B10, doi:10.1029/2008JE003087.
1. Introduction [2] Most of Venus’s surface consists of gently rolling plains with little relief. There are several depressions and two large highlands, Ishtar Terra in the northern hemisphere and Aphrodite Terra along the equator. The Magellan radar images indicated a violent volcanic activity in the planet’s past [Saunders et al., 1992; Solomon et al., 1992]. Extended lava plains, dotted with isolated shield volcanoes and volcanic constructs dominate the geology. The deformations on Venus are not a result of Earth like plate tectonic. They are probably related to the dynamic forces within the planet’s mantle [Basilevsky and Head, 1988, 2002]. [3] Little is known about the surface material. Most of it appears to consist of basalt, but the primary material is poorly classified. The volcanic landforms are consistent with low-viscosity eruptions, which are characteristic of mafic materials like basalt [Head et al., 1992]. However, 1 Institut fu¨r Planetologie, Westfa¨lische Wilhelms-Universita¨t Mu¨nster, Mu¨nster, Germany. 2 Institute of Planetary Research, German Aerospace Center, Berlin, Germany. 3 Department of Marine Remote Sensing, Remote Sensing Technology Institute, German Aerospace Center, Berlin, Germany. 4 LESIA, Observatoire de Paris, CNRS, UPMC, Universite´ ParisDiderot, Meudon, France. 5 Instituto di Astrofisica Spaziale e Fisica Cosmica, Rome, Italy.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003087$09.00
Magellan data suggest that some high-viscosity lava formed pancake domes and festoons [McKenzie et al., 1992; Taylor, 2006], which could be related to the presence of more felsic materials like rhyolite. [4] Recent estimates have demonstrated that there may be significant spatial variations in the surface emissivity as large as 20%, which correspond to the difference between granitic and basaltic rocks [Hashimoto and Sugita, 2003]. The chemical and mechanical erosion of the primary material is controlled by reactions of the surface with the hot and dense atmosphere. The study of the chemical composition and the physical properties of the lower atmosphere and surface of Venus are crucial for the understanding of the surface evolution and the climate of the planet. [5] Although Venus Express primarily focuses on atmospheric science, it also supplies valuable information about the surface. VIRTIS, the Visible and Infrared Thermal Imaging Spectrometer [Drossart et al., 2007; Piccioni et al., 2007], is one of the important experiments dedicated to atmospheric and surface data extraction. VIRTIS is the first instrument operating in an orbit around Venus with the capability to systematically investigate the nightside emission of the planet in the near-infrared atmospheric windows. It performs the first detailed global exploration of the depths of the thick Venusian atmosphere. It provides for the first time clues to the emissivity of surface materials and it may provide direct evidence of active volcanism if present [Baines et al., 2006]. The feasibility of such studies was demonstrated first by Galileo/NIMS [Carlson et al., 1991]
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Figure 1. The three orbit groups with different topographical variations along their footprints. Group 1 orbits (left to right) are 82, 86, 90, 92, 94, 97, 100. Group 2 orbits are 110, 113, 116. Group 3 orbits are 146– 154. The colored background is the surface topography as obtained by Magellan. Each spatial pixel footprint center corresponds to a white dot. The pixel density is higher near the equator because the spacecraft is slower there with respect to the surface. The black edging is used to stress the footprints. Magellan data are available at http://pds-geosciences.wustl.edu/missions/magellan. and Cassini/VIMS [Baines et al., 2000] during the Venus flybys. [6] VIRTIS measurements of the thermal emission from the Venus nightside reveal a high variability of near-infrared spectra over the planet. This is mainly a consequence of spatial variations of cloud opacity and surface elevation, but differences are also due to changes in atmospheric temperature and absorber contents as well as variations of the surface emissivity. While the maximum information on surface and deep atmosphere-surface interaction is obtained from the spectral windows located between 1.00 and 1.35 mm, the windows at 1.74 and 2.3 mm provide information about atmospheric temperature and composition below the main cloud deck. [7] We focus our work on the extraction of surface data from VIRTIS measurements. The present approach is based on a data selection from orbits over the northern hemisphere where the footprints cover a broad range of surface elevation variations including deep valleys like Guinevere Planitia (50°N, 290°E) and high mountain regions like Ishtar Terra (70°N, 340°E) and Beta Regio (30°N, 280°E). A quantitative evaluation of the measurements requires detailed radiative transfer simulations that include appropriate spectral line databases, deep atmosphere continuum absorption features, and multiple scattering effects due to the dense cloud deck. The main intention of part I of the paper, however, is a feasibility study and an estimation of a
quantitative approach to extracting surface information instead of elaborated quantitative retrievals.
2. Data Selection [8] The data were selected from VIRTIS-M-IR measurements over the northern hemisphere of Venus. According to Figure 1, each narrow stripe extends roughly from the equator to the North Pole. It is part of the measurements obtained during one Venus Express (VEX) orbit. One stripe consists of a series of exposures during one VEX orbit. Each exposure yields a frame of 64 spatial pixels times 432 spectral pixels. The 64 pixel wide swath corresponds to a width of roughly two degrees of longitude in the surface footprint, while the latitude remains approximately constant for any pixel in the frame. Thus, the coverage of the pixel footprints of the selected data resembles a latitudinal cross section of spectra at a fixed longitude. Because of the highly elliptical shape of the orbits, a pixel footprint on the surface is of the order of approximately 4 km in longitudinal and 1 km in latitudinal direction at 20°N, but only 0.4 km times 0.1 km at 83°N. [9] The required stripes were found by searching for pushbroom observations with small observation angles close to nadir. This ensures a minimal atmospheric influence on the measured signatures. In this way, the viewing geometry is approximately constant within one stripe, except for a framewise progression of the latitude of the
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Figure 2. Initial and stray light-corrected nightside spectrum of VIRTIS (orbit 149, 9°S). The main NIR window positions are 1, 1.02 mm (9765 cm1); 2, 1.10 mm (9105 cm1); 3, 1.18 mm (8465 cm1); 4, 1.28 mm (7843 cm1); 5, 1.31 mm (7657 cm1); 6, 1.74 mm (5770 cm1); 7, 2.30 mm (4357 cm1). pixel footprints. Furthermore, only measurements at solar incidence angles exceeding 98 degrees were considered for nightside emission studies to ensure that they had been recorded at an adequate distance from the scattered sunlight close to the terminator. [10] To simplify matters, the stripes are named for the corresponding orbit number. Considering the above defined search criteria, three groups with different topographical variations along each group were selected on the northern hemisphere, as shown in Figure 1, to investigate the influence of surface elevation on the spectra. Group 1 contains the orbits 82, 86, 90, 92, 94, 97, and 100, where the topographical variation with latitude is fairly small. These orbits cover lowlands from 10°S to 80°N and 237°E to 265°E. Stronger topographical variations with latitude occur in groups 2 and 3, respectively. Group 2 comprises orbits 110, 113, and 116. The footprints range from 5°S to 83°N and 270°E to 288°E. The surface elevation rises up to 5600 m at Beta Regio. It goes down to 1.6 km at Guinevere Planitia and it is fairly uniform in the north. Group 3 contains orbits with similar topographical variations with latitude as in group 2, but these orbits, with the numbers 146 to 154, span high mountains in high latitudes (Ishtar Terra region at an elevation of up to 6000 m) and lowlands down to 1.5 km elsewhere. They range from 10°S to 83°N and 332°E to 345°E. [11] Some of the selected spectra show a strong increase of radiance with decreasing wavelength in the 1 – 2.5 mm range; that is, the lower envelope in that spectral range is lifted as it is illustrated by the dashed line in Figure 2. A scrutiny of the viewing geometry reveals that this effect is due to direct sunlight close to, but outside the VIRTIS field of view, although there is dark night on the planet’s surface. An examination of the most extreme examples (orbits 137 and 139) shows that the spectral characteristic changes significantly over a small footprint distance and time interval. The change is correlated with the exposure point, as similar occurrences along other orbits confirm. When the Sun angle (i.e., the angle between observation direction and the instrument-to-Sun vector) decreases, the stray light
initially increases, but it disappears abruptly at the point where the Sun becomes eclipsed by Venus itself from the VIRTIS point of view. This stray light effect must not be confounded with upper atmosphere scattering close to the terminator [Meadows and Crisp, 1996] and with indirect light from the illuminated crescent of Venus scattered into the VIRTIS field of view [Mu¨ller et al., 2008], respectively. It is not a surface or atmospheric feature, but is due to direct sunlight striking the VIRTIS instrument. [12] The orbits 137 and 139 with long (3.3 s) and short (0.36 s) exposure times, respectively, can be used to remove the stray light effect to a large extent, because it is most pronounced here and vanishes quickly upon entering eclipse. The difference in radiation measured immediately before and after the eclipse is a good indicator of the amount of stray light. The correction algorithm averages the difference over all pixels of the corresponding frames to ensure a good statistical relevance. Approximately 1 min elapses between the two measurements. This corresponds to a footprint distance of just 125 km. Using the difference between the two measurements seems to be justified, because it will rarely be possible to identify real surface or atmospheric parameter variations in the measured signals along that distance, since the predicted IR surface resolution limit is about 100 km due to the strong cloud scattering [Moroz, 2002a]. [13] The resulting correction curves for the two exposure times are scaled for each individual spectrum to fit it as a lower envelope. This correction is applied to all the orbits and spectra investigated in this work, regardless of the degree of direct sunlight illumination. This prevents a possible discontinuity between spectra that have, or have not, been corrected. The solid line in Figure 2 displays the same spectrum as the dashed line but with the effect of scattered sunlight being removed. Some low-level radiance discontinuities were also eliminated. [14] Because of the limited surface resolution mentioned above it is reasonable to enhance the statistical stability by averaging the spectra of adjacent pixels, where atmospheric conditions are assumed to be similar. The basic data of
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Figure 3. Variability of VIRTIS spectra (orbit 151) at different northern latitudes. Window signatures at 2.3 and 1.7 mm are sometimes very small and may even disappear. every orbit are confined to less than two degrees in longitude. The orbits are divided into bins of one degree in latitude. The mean over all pixels of each bin is then written as a function of latitude. This procedure is applied to the radiance values as well as to the corresponding Magellan elevations. [15] Figure 3 presents a set of four VIRTIS-M spectra in the wavelength range from 1.0 to 2.6 mm taken along orbit 151 (group 3) at different latitudes over the northern hemisphere. This orbit crossed the equatorial lowlands as well as high mountain regions at Ishtar Terra. Figure 3 shows that VIRTIS measurements of the thermal emission from the Venus nightside reveal a high variability of nearinfrared spectra over the planet. The main reasons for the observed differences are spatial variations of cloud optical depth and surface elevation, but their influence can be very clearly distinguished in dependence on wavelength as will be shown below. Window signatures at 2.3 and 1.7 mm are sometimes very small and may even disappear. This is due to an extreme cloud thickness in these cases. Radiation changes may be also due to changes in atmospheric temperature profile and absorber contents and also due to variations in the surface emissivity, but their influence is usually much smaller.
3. Preliminary Radiative Transfer Simulations [16] A radiative transfer model is used to simulate VIRTIS-M spectra. It allows for absorption, emission, and multiple scattering by atmospheric gaseous and particulate constituents in planetary atmospheres. The algorithm was described in detail by Haus and Titov [2000], who applied it to the atmosphere of Mars. Although Earth, Mars, and Venus differ substantially with respect to their total atmospheric thermal regime and composition, the basic method of radiative transfer simulation is the same. It can be applied to quantitative analyses even in the field of environmental and air pollution research [Haus et al., 1994, 1998]. The atmosphere of Venus above the dense cloud deck is rather similar to the conditions in the Earth’s stratosphere and mesosphere with respect to its thermal regime. Therefore, only minor
modifications are required in the radiative transfer algorithm to simulate VIRTIS dayside measurements. The extreme temperature and pressure regime in the deep atmosphere of Venus, however, makes high demands on new theoretical and methodical work including hot gaseous absorption bands, far spectral line wings and pressure-induced absorption. Much effort has been devoted to explaining the effects of spectral line shape under high temperature and pressure conditions [e.g., Burch et al., 1969; Tonkov et al., 1996; Filippov and Tonkov, 1998; Ma et al., 1999; Tvorogov and Rodimova, 1995; Afanasenko and Rodin, 2005], but up to now, the related works do not offer a self-sufficient model that is suitable for practical calculations over the full spectral range that is of interest for surface and deep atmosphere studies. This is the reason why currently most simulation algorithms make use of empirically determined continuum opacities to fit the observations. [17] Unless satisfactory models or laboratory measurements are available, the measured VIRTIS radiation spectra are used to determine the relative magnitude of continuum absorption as well as cloud absorption and scattering features from the signatures themselves. Therefore, the radiative transfer simulation that has been applied throughout this work is a preliminary one, but it will be further developed in near future. The main intention of this paper is to provide a feasibility study and an estimation of a quantitative approach to extracting surface information from VIRTIS data instead of elaborated quantitative retrievals. Nevertheless, the simulation results obtained so far look very promising as is shown below. [18] Look up tables of quasi-monochromatic absorption cross sections k (n, p, T) of gaseous constituents in the atmosphere of Venus are calculated on the basis of a lineby-line procedure, where n is the wave number [cm1], p the atmospheric pressure [mbar], and T the temperature [K] covering the altitude range 0 – 140 km. Cross sections of SO2, HF, HCl, and OCS are evaluated based on the HITRAN 2004 catalog [Rothman et al., 2005] where a Voigt profile with a line cut of 125 cm1 is used. H2O and CO cross-section calculations (line cut 125 cm1) use the HITEMP 1995 catalog [Rothman et al., 1995]. Since this
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Figure 4. Radiative transfer simulation results in a purely gaseous Venus atmosphere. Top three lines are pure CO2 for different spectral line databases; line with diamonds is H2O added to CO2-CDSD, bottom curve (triangle dn) is with continuum included. H2O mixing ratio is 25 ppm (0 – 46 km), Voigt line profiles, line cut 125 cm1. older database only contains H2O lines of the main isotope 161, the minor isotopic species were added from HITRAN04. The HDO abundance is considered to be 150 times the telluric one [deBergh et al., 1991]. The even higher abundance in the upper atmosphere [Bertaux et al., 2007] is not considered at present. HITRAN04 and HITEMP95 were found to be inadequate for simulating near-infrared emissions from Venus nightside with respect to the main constituent, carbon dioxide. Carbon dioxide opacities were finally calculated by using line parameters from the high-temperature carbon dioxide spectroscopic databank (CDSD) compiled by Tashkun et al. [2003]. Figure 4 compares the results of a pure CO2 radiative transfer simulation in the spectral range 1.0– 1.45 mm when the three different databases are used. The high number of additional week spectral lines that are contained in the CDSD database yields many additional absorption features under the high temperature and pressure conditions in the deep Venusian atmosphere. Significant changes in comparison with HITRAN04 and HITEMP95 also occur at longer wavelengths. [19] The diamond line in Figure 4 marks the strong influence of water vapor on the 1.10 and 1.18 mm windows and the intermediate range. It is evident that the striking 1.10 mm window feature only appears when H2O is considered in the simulation. A constant volume mixing ratio of 25 ppm from 0 to 46 km altitude was used. A comparison of this pure gas simulation with VIRTIS measurements (compare Figures 2 and 3) reveals a disagreement of spectral profiles in the range between the 1.10 and 1.18 mm windows. It is known that water vapor spectral lines exhibit a collision-induced superLorentzian behavior leading to higher wing absorptions compared to a pure Lorentz (or Voigt) profile. This effect was not considered here. Contrary to water vapor, carbon dioxide lines are rather characterized by a sub-Lorentz profile, which provides less absorption in the distant line wings. The calculated absorption cross sections k (n 0, p, T) also strongly depend on the line cut criterion that is
applied to suppress contributions from very distant lines relative to the wave number n 0. Different sub-Lorentz line profiles of the simple form fL(n n 0)exp[a(jn n 0j s)] for jn n 0j > s as well as more sophisticated profiles from the literature [Fukabori et al., 1986; Perrin and Hartmann, 1989; Pollack et al., 1993; Meadows and Crisp, 1996; Tonkov et al., 1996], combined with different line cut assumptions (125 –1000 cm1), are currently under investigation with respect to individual band positions of CO2, but a final definition of a ‘‘best choice model’’ could not be provided up to now. As was mentioned above, pressure-induced absorption is currently not included in the simulations. [20] Empirical continuum absorption coefficients have been used so far in window ranges where it is needed. The continuum cross section kc is currently determined from a Venus ‘‘reference spectrum’’ (compare Figure 5) when the clouds are included in the simulation procedure. The bottom line (inverted triangle) in Figure 4 shows the influence of continuum absorption on a pure gas simulation in the 1.0– 1.45 mm range. With the continuum included, the simulated relative peak values and profiles of the window signatures are much closer to the observed ones. [21] Mie scattering theory is applied to derive the microphysical parameters of the H2SO4 clouds, where a four modal log normal size distribution is used (modes 1, 2, 20, and 3 according to Pollack et al. [1993]). Optical depths of cloud absorption, scattering and extinction are calculated on the basis of a particle number density model that essentially corresponds to that described by Moroz [2002b]. The total cloud opacity of a so-called standard model, which is applied throughout this paper, is 26.9 at 1 mm. Unity total optical depth occurs near the upper limit of the main cloud deck at an altitude of 70 km. There are many open questions with respect to the chemical composition and the particle size distribution of the clouds [Moroz, 2002b; Taylor, 2006]. It is well established that the upper cloud particles (mainly mode 2 and mode 20) consist of a water solution of sulphuric acid (75% H2SO4 by weight). The composition of the
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Figure 5. Comparison of VIRTIS measurement for orbit 90, 30°N (solid line) and radiative transfer simulation (dashed line). The surface elevation is 1.0 km. The integers refer to the band number introduced in Figure 2. This spectrum was used as reference spectrum for the simulation model development. smaller mode 1 droplets is unknown. The large mode 3 particles are probably highly variable in composition and spatial and temporal distribution. [22] Rayleigh scattering optical depths are considered according to Hansen and Travis [1974]. The total Rayleigh optical depth at 1 mm is 1.54; that is, it is not negligible within the short-wavelength windows. Radiation multiple scattering in the dense cloudy atmosphere is considered by a successive order procedure where the source functions are calculated in a two-stream approximation of the angulardependent radiation field [Arnold et al., 2000; Haus and Titov, 2000]. [23] The synthetic quasi-monochromatic intensity spectra at the model top level of the atmosphere (140 km) that are generated by the radiative transfer code are convolved with the VIRTIS spectral response function. A constant spectral resolution of 10 nm and a Gaussian instrumental response function have been used so far. In the wave number frame, 10 nm correspond to 100 cm1 at 10000 cm1 (1 mm), 25 cm1 at 5000 cm1 (2 mm), and 4 cm1 at 2000 cm1 (5 mm), respectively. [24] Figure 5 compares a synthetic spectrum with the VIRTIS measurement performed at 30°N on orbit 90. The surface elevation according to Magellan topography is 1.04 km there. This spectrum was selected as the ‘‘reference spectrum’’ to investigate the combined effects of cloud scattering and gas continuum absorption. The spectral range from 1 to 2.6 mm shown in Figure 5 covers the atmospheric windows located at 1.02, 1.10, 1.18, 1.28, 1.31, 1.74, and 2.3 mm. The simulation was performed for the Venus standard atmosphere [Seiff et al., 1985]. An update of these VIRA data by VIRA-2 [Moroz and Zasova, 1997] was not required for the investigations intended here, since the temperature and pressure profiles below 50 km remained unchanged. The calculations indicate that temperature changes within and above the clouds do not influence the surface and deep atmosphere signals. [25] The preliminary quantitative simulation approach yields a more or less good fit to the observed spectrum,
although there are some discrepancies (mainly at 1.7 and 2.3 mm). They are probably due to an inappropriate choice of spectral line profiles, line cut and continuum parameters. Additionally, the lower atmospheric temperature profile may be slightly different from the VIRA standard profile. Small changes can result in significant window flank shifts as corresponding calculations have shown. This will be investigated in more detail in near future. The present gas simulation uses Voigt line shapes and a line cut of 125 cm1 everywhere for each molecular constituent. Pressure-induced continuum absorption is assumed to depend on the total molecular density squared. It is determined as a correction value that must be applied to fit the measured reference spectrum when the total aerosol and Rayleigh optical depths of the atmosphere are 26.9 and 1.54 at 1 mm, respectively. The following values (in units of [cm1 amagat2]) were used in the simulation shown in Figure 5: windows 3 – 5 (1.18, 1.28, 1.31 mm), 1.56 1009; window 6 (1.74 mm), 6.23 1009; window 7 (2.30 mm), 4.17 1008. Pollack et al. [1993] derived the values 7.00 1009 and 2.50 1008 for windows 6 and 7, respectively, while Marcq et al. [2006] used a value of 3.5 108 at 2.3 mm. The water vapor signatures, which appear in the pure gas simulation between 1.10 and 1.18 mm (compare Figure 4) due to the neglect of super-Lorentz line shapes, have been currently removed by applying a continuum in the order of 1.0 108 at 1.13 mm. All the continuum absorption coefficients are assumed to be independent of wavelength throughout a given window range. They are also not varied from spectrum to spectrum; that is, the values are held constant for simulations with respect to different VIRTIS measurements.
4. Results [26] Figure 6 illustrates the simulated radiances at the top level of the Venus atmosphere for cloud optical depths (opacities) that differ from the standard model by a factor of 0.7, 1.0, 1.4, 2.0, and 2.5, respectively. The surface
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Figure 6. Simulated nightside radiance spectra in dependence on cloud opacity factor. Radiances decrease with increasing factor. Factor 1 corresponds to a total cloud optical depth of 26.9 at 1 mm. The surface elevation is 1.0 km. elevation is constant at a value of 1.04 km. Each window is sensitive to changes in the cloud opacity, but the contrasts increase toward longer wavelengths. Cloud contrast is most apparent within the 2.3 mm window. The aerosol single scattering albedo below 1.5 mm is very close to unity; that is, conservative scattering is the dominating cloud feature. In other words, the clouds do not contribute to thermal emission, but act as a strong attenuator of radiation that originates from below the cloud deck and emerges into space. Little spectral dependence of upwelling radiation is added by the clouds below 1.5 mm. This is the reason why it should be possible to use multispectral radiation measurements of the Venus surface and lower atmosphere in the 1.0 – 1.35 mm window complex to eliminate cloud interferences. [27] Figure 7 illustrates the change of averaged window radiances in dependence on the cloud opacity factor. Averaging means that mean values over each window range are
used instead of the maximum value. The ranges [mm] were defined as follows. W1, 1.028– 1.035; W2, 1.090 – 1.110; W3, 1.160 – 1.190; W4, 1.255 – 1.289; W5, 1.300 – 1.309; W6, 1.691 – 1.780; W7, 2.200 – 2.290. Only the shortwavelength part of window 7 at 2.3 mm was used here to eliminate possible minor constituent variations (H2O, CO, COS, SO2, HF) in the long-wavelength part at the right. The left-hand part is exclusively determined by CO2. Water vapor concentration changes, however, which may affect the windows at 1.74, 1.18, and 1.10 mm, respectively, were not eliminated at present. [28] As expected, there is a dependence of radiation on cloud opacity in the windows 1– 5 that leads to constant radiance ratios. There is an increasing contribution of cloud absorption with increasing wavelength. Thus, the linear dependence is lost and the radiances at 1.7 and 2.3 mm decrease much faster with growing cloud opacity. As a
Figure 7. Simulated averaged window radiances as a function of cloud opacity factor. The integers in the box refer to the band number introduced in Figure 2. The surface elevation is 1.0 km. 7 of 13
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Figure 8. Comparison of measured (scatter points) and simulated (lines) averaged window radiance ratios as a function of 2.3 mm averaged window radiance. The scatterplot of binned measured ratios includes all latitudes from the 19 orbits of groups 1 – 3 (compare Figure 1), where the surface elevation is constrained to the interval [1500 m, 500 m]. The integers in the box refer to the band numbers introduced in Figure 2. Values included in parentheses denote the corresponding nominal wavelengths. F is the cloud opacity factor. Nearly constant ratios for the surface windows are apparent. consequence, a cloud correction algorithm, which makes use of radiance ratios, tends to fail for longer wavelengths. [29] Figure 8 compares measured and simulated averaged window radiance ratios as a function of the wavelengthaveraged radiance that was measured at 2.3 mm. The radiance at 2.3 mm is most sensitive to changes in cloud optical depth as is clearly visible in Figures 6 and 7, and also in Figure 3. It is insensitive to changes in surface elevation (compare Figures 9 and 10). Cloud opacity is not a free parameter in the measurements as it is in the simulations. It cannot be measured directly by the VIRTIS instrument. But the observed strong variability at the lefthand flank of the 2.3 mm window is a good indicator of cloud thickness on the Venus nightside. Thus, a scatterplot of measured radiance ratios against the measured 2.3 mm
radiance, which decreases with increasing cloud opacity, has the same physical background as a plot of simulated ratios against the cloud opacity. This is convincingly demonstrated in Figure 8. The scatterplot is restricted to measurements from the 19 orbits shown in Figure 1, and the surface elevation according to Magellan topography is constrained to the interval [1500 m, 500 m]. The restriction to a narrow elevation interval is necessary to omit strong radiation changes due to topographical changes (compare Figure 9). It is assumed that the quantitative differences between measurement plot and simulation result, which are visible in the case of 6/1 and 7/1 ratios, are due to present inaccuracies in the spectral line profiles and continuum coefficients at these wavelengths.
Figure 9. Simulated nightside radiance spectra in dependence on surface elevation for the standard cloud model (opacity factor 1.0). Radiances decrease with increasing elevation in the surface windows. The 1.74 and 2.3 mm windows are unaffected. 8 of 13
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Figure 10. Simulated averaged window radiances as a function of surface elevation. The integers in the box refer to the band numbers introduced in Figure 2. The cloud opacity factor is 1.0. [30] Another important conclusion can be derived from Figure 8. All the measurement points are distributed up to a maximum 2.3 mm radiance value of about 0.065 W/(m2 sr mm). This value corresponds to a cloud opacity factor of 0.95 (compare Figure 7) and seems to indicate that the assumed standard cloud optical depth of 26.9 (opacity factor 1.0) is rather the lower limit of what is present in the real Venusian atmosphere. Since continuum parameters were derived based on this standard model, a possible conclusion is that not only higher cloud opacities, but also lower continuum values are closer to reality than those adopted in the present ‘‘standard case model.’’ This will be addressed in future investigation. [31] The possibility to eliminate cloud signals from the VIRTIS measurements below 1.5 mm by the use of radiance ratios is of decisive importance for our future work that will mainly focus on Venus surface data extraction and the study of surface and lower atmosphere interactions. [32] Figures 9 and 10 illustrate the change of simulated radiances with surface elevation and wavelength. The cloud opacity was held constant in the simulations. The windows at 1.02, 1.10, and 1.18 mm exhibit a clear dependence of transmitted radiation on topographical features and, thus, on surface thermal emission. The radiance decreases considerably with increasing elevation, since high-elevation surfaces are substantially cooler and emit less thermal radiation. An elevation change of 12 km results in a surface temperature change of almost 100 K. The window at 1.28 mm still shows a minor influence of surface temperature, while it is mainly determined by features of the lower atmosphere (15 – 30 km). The surface elevation (temperature) does not observably influence the 1.31, 1.74, and 2.3 mm windows radiances, respectively. A simulation was performed, where the surface contribution was excluded assuming a surface emissivity value of zero. Thus, a quantitative estimate of the surface thermal emission contribution to emission within the Venus atmospheric windows was obtained. In the windows 1 (1.02 mm), 2 (1.10 mm), 3 (1.18 mm), and 4 (1.28 mm), the surface contributes 97 (95) %, 59 (60) %, 27 (40) %, and 1.3 (<2) %
to the total emission. The values in parentheses are the corresponding results of Meadows and Crisp [1996]. [33] Figure 11 compares measured and simulated averaged window radiance ratios as a function of surface elevation. Since the radiance at 1.02 mm is most sensitive against elevation changes, the ratios are not constant and, thus, a valuable measure of surface topography and surface temperature, respectively. The scatterplot contains all measurement data bins from the 19 orbits shown in Figure 1. As Figure 11 shows, these orbits pass predominantly over lowlands with an elevation between 1 km and 0 km. Another maximum occurs between 2 and 4 km (Beta Regio, Ishtar Terra). There is again a very good qualitative correlation between measurements and simulation.
Figure 11. Comparison of measured (scatter points) and simulated (lines) averaged window radiance ratios as a function of surface elevation. The scatterplot of measured ratios includes all bins from the 19 orbits of groups 1 – 3 (compare Figure 1). The integers in the box refer to the band numbers introduced in Figure 2. Values included in parentheses denote the corresponding nominal wavelengths. The cloud opacity factor in the simulation is 1.0.
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Figure 12. Averaged window radiances at 1.02, 1.18, and 2.30 mm and the radiance ratios 1.18/1.02 in comparison with surface elevation according to Magellan topography as a function of latitude for orbits (a – b) 90, (c– d) 113, and (e – f) 150. The ratios are based on measured radiances, which have been displayed and discussed before by Arnold et al. [2007]. [34] Figures 12a– 12f yield additional evidence of the important finding, that emission window radiance ratios can be used to separate atmospheric and surface properties in the measured Venus spectra. It shows the radiances at 1.10, 1.18, and 2.30 mm and the radiance ratio 1.18/1.02 in comparison with surface elevation according to Magellan topography as a function of latitude for three selected orbits. Orbit 90 (Figures 12a and 12b) at 250°E is representative for group 1 measurements over predominantly low elevation areas in the northern hemisphere of Venus. Orbit 113 (Figures 12c and 12d) touched the center of Beta Regio and the deep depressions of Guinevere Planitia at 283°E, while orbit 150 (Figures 12e and 12f) passed over the highlands of Ishtar Terra at 339°E. Higher elevations should result in
lower radiances of the windows 1 (1.02mm) and 3 (1.18 mm), and vice versa. This is the case at Beta Regio, Guinevere Planitia, and Ishtar Terra in comparison to the surrounding terrains. But it is rather hard or even impossible to verify this behavior for smaller topographical variations. The window 1 and 3 measurements on orbit 90 (Figure 12a), for example, indicate higher radiances at 35°N, but this has nothing to do with an elevation decrease. The high window 7 (2.3 mm) radiance value at this location proves that much lower cloud opacity is the cause. The same holds true on this orbit at 60°N. The ratio 1.18/1.02 (Figure 12b), on the other hand, follows the elevation profile in most details. The elevation subsidiary summit at Ishtar Terra on orbit 150 at 65°N (Figure 12e) is not reflected in decreasing window 1
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Figure 13. Retrieval of surface topography of group 1 and 2 orbits near Beta Regio using 1.18 mm to 1.02 mm radiance ratios. (a) ‘‘The target state’’ of topography retrieval according to Magellan data, which were smoothed over the swaths with respect to the 100 km resolution limit to be comparable with the VIRTIS data. (b) ‘‘The actual state’’ of topography retrieval. The ratio for each radiance measurement on a pixel basis is mapped to the elevation according to the linear fit of measurement data shown in Figure 11 (1.18/1.02 versus Magellan elevation) and depicted in the same color coding as the Magellan elevation. The black edging is used to stress the footprints. and 3 radiances, since it is masked by a local cloud decrease (enhanced window 7 radiance). The ratio 1.18/1.02 (Figure 12f) shows this feature very well. [35] The high correlation between the ratios of radiances measured in the surface windows between 1.0 and 1.35 mm and the planetary topography can be used to extract surface data of Venus. Figure 13 demonstrates this capability for some parts of orbits in the vicinity of Beta Regio including the orbits of groups 1 and 2. Figure 13a shows the topographical ‘‘target state’’ as it is defined by the Magellan data, which were smoothed over the swaths with respect to the 100 km resolution limit to be comparable with the VIRTIS data. Figure 13b is the ‘‘actual state’’ of the topography retrieval from VIRTIS data, which is obtained from a linear regression of the r3/r1 (1.18/1.02) scatterplot shown in Figure 11. The ratio r3/r1 as a function of the Magellan elevation h reads r3/r1(h) = (3.08 ± 0.01) + (0.319 ± 0.008) h/km, with an error of twice the standard deviation. The ratio for each radiance measurement on a pixel basis is mapped to the elevation according to the linear fit of measurement data shown in Figure 11 and depicted in the same color coding as the Magellan elevation. The black edging in Figures 13a and 13b is used to stress the footprints.
[36] In general, the VIRTIS and Magellan topographies are largely in agreement with each other, but differences occur in certain small areas. These areas will be of special interest for future data analyses. The examination of differences at comparable elevations and surface temperatures may give hints either to local variations of the deep atmospheric temperature profiles and composition or to different surface materials. [37] As discussed above, the window 1 at 1.02 mm contributes more than 95% to the total emission of the Venus nightside. Thus, any relative change in the surface emissivity corresponds to approximately the same relative radiance change. It should be possible to detect a difference of a few percent. According to Hashimoto and Sugita [2003], there may be spatial variations in the surface emissivity of Venus as large as 20%, which correspond to the difference between granitic (rhyolitic) and basaltic rocks due to a different iron oxide content in the rocks.
5. Conclusions [38] The VIRTIS-M nightside IR data form a valuable basis for systematic and continuous surface and deep atmosphere studies of Venus. The observed high variability
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of measured signatures is mainly due to spatial variations of cloud optical depth and surface elevation. The conservative character of cloud scattering below 1.5 mm, where the clouds do not contribute to thermal emission, makes it possible to use multispectral radiation measurements of the Venus surface and lower atmosphere emission in the 1.0– 1.35 mm window complex to separate cloud interference and surface influences. [39] The radiance in the left-hand flank of the 2.3 mm window does not depend on surface features and minor constituent distribution changes in the lower atmosphere. It is, therefore, a good indicator of cloud thickness on the Venus nightside and can be used for quantitative estimates. From the comparison of measured and simulated averaged radiance ratios as a function of 2.3 mm averaged radiance and, hence, of cloud opacity, it can be concluded that a cloud optical depth in the order of 27 is a lower limit with respect to the VIRTIS data that have been investigated so far. [40] Measurements as well as radiative transfer simulations have proven a high correlation between the radiance ratios in the emission windows between 1.0 and 1.35 mm and the surface elevation. This allows a mapping of surface topography and the retrieval of surface temperature. On the basis of measured radiances at 1.18 and 1.02 mm and their ratios, the surface topography in selected regions around Beta Regio has been retrieved and compared with the Magellan data. The way the results largely agree proves that the ratio method to decloud the VIRTIS data is well suited to extract surface information. Differences have been observed in certain small areas. To the first order, the radiance ratios are a function of surface temperature. Small deviations from this first-order dependence are possibly due to surface materials that exhibit different emissivity characteristics. Mu¨ller et al. [2008] and G. L. Hashimoto et al. (Galileo near-infrared mapping spectrometer data suggests felsic highland crust on Venus, submitted to Journal of Geophysical Research, 2008) deal with observed emissivity variations. [41] The preliminary radiative transfer simulations have demonstrated the capability of the algorithm to investigate the surface of Venus. Future studies will focus on a separation of atmospheric and surface radiance contributions in the measured VIRTIS signals and a possible separation of surface temperature and emissivity by a comparative investigation of different surface windows and their radiance ratios. Since the orbit repetition increases during the mission operation, it will be possible to distinguish between static and dynamic local variations and to gather more detailed information about the surface of Venus on a global scale. [42] These objectives can only be achieved by further progress in the radiative transfer simulations to eliminate the masking of the Venus nightside windows by far wing and pressure-induced absorptions of the deep atmosphere constituents. A great deal of theoretical and numerical effort will be required, including taking appropriate spectral line databases and line profiles into account, as well as revised models of cloud composition. [43] Acknowledgments. We acknowledge the work of the entire Venus Express team which allowed these data to be obtained.
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References Afanasenko, T. S., and A. V. Rodin (2005), The effect of collisional line broadening on the spectrum and fluxes of thermal radiation in the lower atmosphere of Venus, Sol. Syst. Res., 39(3), 187 – 198, doi:10.1007/ s11208-005-0034-1. Arnold, G., R. Haus, and H. Hirsch (2000), The Planetary Fourier Spectrometer (PFS): Studies of the Martian Atmosphere and Surface, Res. Rep. 2000-18, 114 pp., German Aerosp. Cent., Inst. fu¨r Weltraumsensorik und Planetenerkundung, Berlin. Arnold, G., W. Do¨hler, R. Haus, D. Kappel, P. Drossart, G. Piccioni, and the VIRTIS Team (2007), Estimation of a quantitative approach for Venus surface data extraction from VIRTIS measurements using topographical variations, Geophys. Res. Abstr., 9, 07972. Baines, K. H., et al. (2000), Detection of sub-micron radiation from the surface of Venus by Cassini/VIMS, Icarus, 148, 307 – 311, doi:10.1006/ icar.2000.6519. Baines, K. H., S. Atreya, R. W. Carlson, D. Crisp, P. Drossart, V. Formisano, S. S. Limaye, W. J. Markiewicz, and G. Piccioni (2006), To the depths of Venus: Exploring the deep atmosphere and surface of our sister world with Venus Express, Planet. Space Sci., 54, 1263 – 1278, doi:10.1016/j.pss.2006.04.034. Basilevsky, A. T., and J. W. Head III (1988), The geology of Venus, Annu. Rev. Earth Planet. Sci., 16, 295 – 317, doi:10.1146/annurev.ea. 16.050188.001455. Basilevsky, A. T., and J. W. Head III (2002), Venus: Timing and rates of geologic activity, Geology, 30, 1015 – 1018, doi:10.1130/00917613(2002)030<1015:VTAROG>2.0.CO;2. Bertaux, J. L., et al. (2007), A warm layer in Venus’ cryosphere and highaltitude measurements of HF, HCl, H2O and HDO, Nature, 450, 469 – 646, doi:10.1038/nature05974. Burch, D. E., D. A. Gryvnak, R. R. Patty, and C. E. Bartky (1969), Absorption of infrared radiant energy by CO2 and H2O. IV. Shapes of collision broadened lines, J. Opt. Soc. Am., 59, 267 – 280. Carlson, R. W., K. H. Baines, L. W. Kamp, P. R. Weissman, W. D. Smythe, A. C. Ocampo, T. V. Johnson, D. L. Matson, J. B. Pollack, and D. Grinspoon (1991), Galileo infrared imaging spectroscopy measurements at Venus, Science, 253, 1541 – 1548, doi:10.1126/science.253.5027.1541. deBergh, C., B. Be´zard, T. Owen, D. Crisp, J.-P. Maillard, and B. L. Lutz (1991), Deuterium on Venus: Observations from the Earth, Science, 251, 547 – 549, doi:10.1126/science.251.4993.547. Drossart, P., et al. (2007), A dynamic upper atmosphere of Venus as revealed by VIRTIS on Venus Express, Nature, 450, 641 – 645, doi:10.1038/nature06140. Filippov, N. N., and M. V. Tonkov (1998), Theory of band shapes in molecular spectra of gases: Application to band wings, J. Chem. Phys., 108, 3608 – 3619, doi:10.1063/1.475755. Fukabori, N., T. Nakazawa, and M. Tanaka (1986), Absorption properties of infrared active gases at high pressures—I. CO2, J. Quant. Spectrosc. Radiat. Transfer, 36, 265 – 270, doi:10.1016/0022-4073(86)90074-9. Hansen, J. E., and L. D. Travis (1974), Light scattering in planetary atmospheres, Space Sci. Rev., 16, 527 – 610, doi:10.1007/BF00168069. Hashimoto, G. L., and S. Sugita (2003), On observing the compositional variability of the surface of Venus using nightside near-infrared thermal radiation, J. Geophys. Res., 108(E9), 5109, doi:10.1029/2003JE002082. Haus, R., and D. V. Titov (2000), PFS on Mars Express: Preparing the analysis of infrared spectra to be measured by the planetary Fourier spectrometer, Planet. Space Sci., 48, 1357 – 1376, doi:10.1016/S00320633(00)00116-1. Haus, R., K. Scha¨fer, W. Bautzer, J. Heland, H. Mosebach, H. Bittner, and T. Eisenmann (1994), Mobile FTIS-monitoring of air pollution, Appl. Opt., 33(24), 5682 – 5689. Haus, R., R. Wilkinson, J. Heland, and K. Scha¨fer (1998), Remote sensing of gas emissions on natural gas flares, Pure Appl. Opt., 7(4), 853 – 862, doi:10.1088/0963-9659/7/4/020. Head, J. W., L. S. Crumpler, J. C. Aubele, J. E. Guest, and R. S. Saunders (1992), Venus volcanism: Classification of volcanic features and structures, associations, and global distribution from Magellan data, J. Geophys. Res., 97, 13,153 – 13,198. Ma, Q., R. H. Tipping, C. Boulet, and J.-P. Bouanich (1999), Theoretical far-wing line shape and absorption for high-temperature CO2, Appl. Opt., 38, 599 – 604, doi:10.1364/AO.38.000599. Marcq, E., T. Encrenaz, B. Be´zard, and M. Birlan (2006), Remote sensing of Venus’ lower atmosphere from ground-based IR spectroscopy: Latitudinal and vertical distribution of minor species, Planet. Space Sci., 54, 1360 – 1370, doi:10.1016/j.pss.2006.04.024. McKenzie, D., P. G. Ford, F. Liu, and G. H. Pettengill (1992), Pancakelike domes on Venus, J. Geophys. Res., 97, 15,967 – 15,976. Meadows, V. S., and D. Crisp (1996), Ground-based near-infrared observations of the Venus nightside: The thermal structure and water abundance
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near the surface, J. Geophys. Res., 101(E2), 4595 – 4622, doi:10.1029/ 95JE03567. Moroz, V. I. (2002a), Estimates of visibility of the surface of Venus from decent probes and balloons, Planet. Space Sci., 50, 287 – 297, doi:10.1016/S0032-0633(01)00128-3. Moroz, V. I. (2002b), Studies of the atmosphere of Venus by means of spacecraft: Solved and unsolved problems, Adv. Space Res., 29, 215 – 225, doi:10.1016/S0273-1177(01)00571-3. Moroz, V. I., and L. V. Zasova (1997), VIRA-2: A review of inputs for updating the Venus International Atmosphere, Adv. Space Res., 19, 1191 – 1201. Mu¨ller, N., J. Helbert, G. L. Hashimoto, C. C. C. Tsang, S. Erard, G. Piccioni, and P. Drossart (2008), Venus surface thermal emission at one micrometer in VIRTIS imaging observations: Evidence for variation of crust and mantle differentiation conditions, J. Geophys. Res., doi:10.1029/2008JE003118, in press. Perrin, M. Y., and J. M. Hartmann (1989), Temperature-dependent measurements and modeling of absorption by CO2-N2 mixtures in the far linewings of the 4.3 mm CO2 band, J. Quant. Spectrosc. Radiat. Transfer, 42, 311 – 317. Piccioni, G., et al. (2007), South polar features on Venus similar to those near the north pole, Nature, 450, 637 – 640, doi:10.1038/nature06209. Pollack, J. B., et al. (1993), Near-infrared light from Venus’ nightside: A spectroscopic analysis, Icarus, 103, 1 – 42, doi:10.1006/icar.1993.1055. Rothman, L. S., R. B. Wattson, R. R. Gamache, J. Schroeder, and A. McCann (1995), HITRAN, HAWKS and HITEMP high-temperature molecular database, Proc. Soc. Photoopt. Instrument. Eng., 2471, 105 – 111. Rothman, L. S., et al. (2005), The HITRAN 2004 molecular spectroscopic database, J. Quant. Spectrosc. Radiat. Transfer, 96, 139 – 204, doi:10.1016/j.jqsrt.2004.10.008. Saunders, R. S., et al. (1992), Magellan mission summary, J. Geophys. Res., 97, 13,067 – 13,090. Seiff, A., J. T. Schofield, A. J. Kliore, F. W. Taylor, S. S. Limaye, H. E. Revercomb, L. A. Sromovsky, V. V. Kerzhanovich, V. I. Moroz, and
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M. Y. Marov (1985), Models of the structure of the atmosphere of Venus from the surface to 100 kilometers altitude, Adv. Space Res., 5, 3 – 58, doi:10.1016/0273-1177(85)90197-8. Solomon, S. C., et al. (1992), Venus tectonics: An overview of Magellan observations, J. Geophys. Res., 97, 13,199 – 13,256, doi:10.1029/ 91JA02410. Tashkun, S. A., V. I. Perevalov, J.-L. Teffo, A. D. Bykov, and N. N. Lavrentieva (2003), CDSD-1000, the high-temperature carbon dioxide spectroscopic databank, J. Quant. Spectrosc. Radiat. Transfer, 82, 165 – 196, doi:10.1016/S0022-4073(03)00152-3. Taylor, F. W. (2006), Venus before Venus Express, Planet. Space Sci., 54, 1249 – 1262, doi:10.1016/j.pss.2006.04.031. Tonkov, M. V., N. N. Filippov, V. V. Bertsev, J. P. Bouanich, N. Van-Thanh, C. Brodbeck, J. M. Hartmann, C. Boulet, F. Thibault, and R. Le Doucen (1996), Measurements and empirical modeling of pure CO2 absorption in the 2.3 mm range at room temperature: Far wings, allowed and collisioninduced bands, Appl. Opt., 35, 4863 – 4870. Tvorogov, S. D., and O. B. Rodimova (1995), Spectral line shape. I. Kinetic equation for arbitrary frequency detunings, J. Chem. Phys., 102, 8736 – 8745, doi:10.1063/1.468977.
G. Arnold and D. Kappel, Institute of Planetary Research, DLR, Rutherfordstrasse 2, D-12489 Berlin, Germany. (
[email protected];
[email protected]) P. Drossart, LESIA, Observatoire de Paris, CNRS, UPMC, Universite´ Paris-Diderot, 5 place Jules Janssen, F-92195 Meudon, France. R. Haus, Department of Marine Remote Sensing, Remote Sensing Technology Institute, DLR, Rutherfordstrasse 2, D-12489 Berlin, Germany. G. Piccioni, INAF-IASF, via del Fosso del Cavaliere 100, I-00133 Rome, Italy.
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114, E00B41, doi:10.1029/2008JE003156, 2009
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Full Article
Venus Express bistatic radar: High-elevation anomalous reflectivity Richard A. Simpson,1 G. Leonard Tyler,1 Bernd Ha¨usler,2 Riccardo Mattei,2 and Martin Pa¨tzold3 Received 31 March 2008; revised 1 April 2009; accepted 17 April 2009; published 27 June 2009.
[1] Magellan (MGN) bistatic radar observations in 1994 confirmed earlier Pioneer
Venus reports of unusual Venus surface reflectivity and emissivity at elevations above 6054 km radius. They also revealed that the anomalous values of surface dielectric constant e near Cleopatra Patera included a large imaginary component (e i 100) at 13 cm wavelength, consistent with a semiconducting surface material. The MGN observations were conducted using a linearly polarized wave, canted at 45° with respect to the plane of incidence and radiated by the MGN synthetic aperture radar antenna toward the specularly reflecting region of the mean planetary surface. In 2006 similar experiments were conducted using 13 cm circularly polarized transmissions from Venus Express (VEX). The VEX signal-to-noise ratio (SNR) was lower than that of MGN, but elevated jej has been inferred broadly over Maxwell Montes. A quasi-specular echo was detected near Cleopatra but with insufficient SNR to address the question of conductivity. An early failure of the VEX 13 cm radio system precludes further measurements with VEX. Citation: Simpson, R. A., G. L. Tyler, B. Ha¨usler, R. Mattei, and M. Pa¨tzold (2009), Venus Express bistatic radar: High-elevation anomalous reflectivity, J. Geophys. Res., 114, E00B41, doi:10.1029/2008JE003156.
1. Introduction [2] Early Earth-based radar images of Venus revealed three very bright areas: ‘‘Alpha’’ and ‘‘Beta’’ near the equator [Goldstein, 1965; Rogers and Ingalls, 1970] and ‘‘Maxwell’’ at 70°N latitude [Jurgens, 1970]. On the basis of then-current understanding, investigators hypothesized that these were regions of very rough surface texture, where highly tilted surface elements preferentially backscattered radar signals toward Earth. On the Moon the large, young crater Tycho had been identified in range-Doppler data as a bright scattering center [Pettengill, 1962]. But the bright features on Venus did not have the appearance of craters, nor did the brightness of Alpha and Beta (both near the equator) vary with the planet’s rotation, as might have been expected had there been a significant specular component to the scattering. [3] The Pioneer Venus Orbiter (PVO) radar altimeter provided the first topographic model for Venus [Pettengill et al., 1980] as well as measurement of the planet’s backscattering parameters and emissivity. Ford and Pettengill [1983] reported that the areas of unusually high reflectivity corresponded to those with low emissivity and that both depended on altitude. With these results, opinion within the planetary community shifted to favor the presence of high 1 Space, Telecommunications, and Radioscience Laboratory, Stanford University, Stanford, California, USA. 2 Institut fu¨r Raumfahrttechnik, Universita¨t der Bundeswehr Mu¨nchen, Neubiberg, Germany. 3 Rheinisches Institut fu¨r Umweltforschung, Cologne, Germany.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003156$09.00
values of surface dielectric constant, e, rather than roughness, as the important surface property underlying the anomalous brightness. The question then became ‘‘What materials could be present on Venus that had e values well above the normal range for terrestrial analogs?’’ Iron pyrite was an early candidate [Ford and Pettengill, 1983]; ferroelectric and other materials have been proposed subsequently [Shepard et al., 1994]. [4] Magellan (MGN) used its side-looking radar, nadir altimeter, and radiometer to map Venus at much higher resolution than PVO, confirming that anomalous reflectivity on Venus occurred globally at radii greater than about 6054 km [Pettengill et al., 1992]. Magellan also showed that the anomalous behavior was less pronounced at the highest elevations, although the thresholds varied somewhat with latitude [Klose et al., 1992]. A unique Magellan contribution was implementation of several bistatic radar (BSR) experiments near the end of its mission. Forwardscattered, quasi-specular reflections from the area around Cleopatra Patera on Maxwell Montes contained a circularly polarized component which could be explained only if the surface material were moderately conducting. For this, Pettengill et al. [1996] proposed a surface composition including elemental tellurium, a trace constituent of terrestrial volcanic gases with a melting point of 723 K. This temperature corresponds to 6054 km elevation in the atmosphere of Venus, above which tellurium can exist as a solid. Schaefer and Fegley [2004] have since shown that elemental tellurium is not stable on Venus and favor certain lead and bismuth compounds. [5] Venus Express (VEX) entered orbit in April 2006, providing a new opportunity for bistatic radar experiments
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E00B41 Table 1. Bistatic Radar Parameters
Magellan Venus Express Venus Express S Band S Band X Band
Parameter
Spacecraft Transmitter Frequency, f (MHz) 2298 Wavelength, l (cm) 13.1 Polarization Linear 5 Power, PT (watts) Antenna diameter (m) 3.7 35.9 Antenna gain, GT (dB) 1.1 Antenna beam widtha (°) Earth Receiver Antenna diameter (m) 70 63.3 Antenna gain, GR (dB) System temperature (K) 22 – 26 Polarization RCP and LCP
2296 13.1 RCP 5 1.3 25.7 3.0
8420 3.6 RCP 65 1.3 37.0 0.8
70 63.3 22 – 30 RCP and LCP
70 74.3 24 – 26 RCP and LCP
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insufficient RCP power to support estimation of jej at incidence angles 8 35°. Bistatic radar experiments were then suspended. This paper summarizes the methods used for analyzing the circularly polarized data and the application to 13 cm data acquired from Maxwell Montes in June 2006.
2. Experiment Overview
a
Antenna beam width is given as the angular distance between the boresight (peak value) and the half power point.
[Ha¨usler et al., 2006]. Magellan’s experiments were conducted from a near-circular, 93 min orbit at both 13 and 3.6 cm wavelengths (Table 1), but only 13 cm echoes were detected from Maxwell Montes. The VEX 13 cm transmitter radiated 5 W, the same power as Magellan but through a smaller antenna, while VEX polarization was the more conventional right circular (RCP). By applying a methodology comparable to determination of the Brewster angle, at which VEX RCP and LCP (left-circularly polarized) echoes should have equal powers, we proposed to carry out additional studies of the high-jej areas found by Magellan. The initial, highly elliptical, 24 h VEX orbit had its periapsis near 70°N latitude, which was favorable for observations of Maxwell. We also anticipated study of other high-elevation targets such as Ovda Regio and Thetis Regio with eight and three times, respectively, the anomalous high-altitude surface area of Maxwell [Pettengill et al., 1992, Table 2]. [6] Nine bistatic experiments were conducted after VEX orbit insertion (Table 2). Five were at incidence angles 8 70° suitable for studying surfaces with large jej. At 8 70° the 13 and 3.6 cm atmospheric absorptions along the propagation path are about 1 and 20 dB, respectively, rendering detection of surface echoes at 3.6 cm impractical with current equipment. A failure aboard the VEX spacecraft in late 2006 reduced the radiated power at 13 cm wavelength by 15 dB. Strong 3.6 cm LCP echoes were seen in the final experiments in August 2007; but there was
[7] The scattering geometry is shown in Figure 1. Transmissions from the VEX spacecraft were directed toward Venus’ surface by the high-gain antenna (HGA). HGA pointing was updated throughout the experiment so that the antenna beam tracked the specular point (8 = 8i = 8r) on the mean spherical surface as the spacecraft moved in its orbit. Because the direct and reflected path lengths changed at different rates, the differential Doppler shift between fd and fr usually allowed easy separation of the two signals. [8] The scattering process is characterized as quasispecular; that is, the total echo comprises many individual reflections from randomly located, specularly oriented, mirror-like facets. There is not a single fr but a spread of frequencies, where each frequency component represents the Doppler shift of one or more individual facets located near the nominal specular point on a smooth planet. The degree of frequency dispersion and the resulting total power received depend on the roughness of the surface weighted by the HGA illumination pattern [Tyler and Ingalls, 1971]. [9] On Earth the echo signals were captured by the 70 m antenna at the NASA Deep Space Network (DSN) site near Canberra, Australia. Signals were amplified, adjusted in frequency to remove the predicted sources of Doppler shifts, converted to baseband, and sampled at the Nyquist frequency (100 kHz) for later processing. Calibration signals using ambient loads and nominal 12.5 K value noise diodes were recorded as part of the data stream to provide references for later scaling to absolute echo powers. [10] During data processing, power spectra hViVi*i were computed for each RCP and LCP sample stream, where h i denotes time averaging, Vi is the baseband complex voltage spectrum from the ith Fourier component, and Vi* is its complex conjugate. Typically, 1024-point spectra were averaged over 10 s (Figure 2). Power in frequencies adjacent to the surface echo was used to establish a noise baseline, which, when compared with the calibration signal output, yielded the absolute received echo signal power.
Table 2. Venus Express Bistatic Radar: Overviewa Date
VEX Orbit Number
Earth Receive Time (UTC)
Earth-Venus Distance (A.U.)
Target
Elevation (km)
Latitude (°N)
Longitude (°E)
Incidence Angle 8 (°)
23 May 2006 15 Jun 2006 17 Jun 2006 19 Jun 2006 18 Mar 2007 20 Mar 2007 12 Jun 2007 2 Aug 2007 3 Aug 2007
32 55 57 59 331 333 417 468 469
0251 0159 0200 0201 0633 0633 0713 0609 0604
1.126 1.281 1.294 1.307 1.305 1.293 0.684 0.325 0.321
‘‘Commissioning’’ Maxwell Montes Maxwell Montes Maxwell Montes Rhea Mons Rhea Mons Ozza Mons Theia Mons Rhea Mons
N/A 6058.7 6058.7 6058.7 6055.5 6055.5 6056.5 6055.8 6055.5
31.3 66.69 67.09 67.44 22.54 24.52 0.15 22.46 21.56
298.3 0.15 3.54 6.97 281.36 285.75 197.99 279.27 280.78
50.83 72.44 72.22 72.00 69.02 70.11 12.95 35.87 35.19
a
Elevations (column 6) are from Table 2 of Pettengill et al. [1992]; targets were chosen for study of anomalous emissivity/reflectivity. The first experiment was conducted to test spacecraft and ground operations; target location was incidental. Earth receive time, latitude, longitude, and incidence angle are nominal values neglecting atmospheric refractivity. Observations typically spanned 1 h, providing profiles several thousands of kilometers in length, including the identified target. Experiments beginning 2007-03-18 were degraded at 13 cm wavelength as the result of a 15 dB loss in spacecraft radiated power (see text).
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Figure 1. Bistatic radar geometry. Transmitter, receiver, and the center of Venus define a plane within which the incidence angle is 8i and reflection angle is 8r. From orbit, Venus Express pointed its high-gain antenna (HGA) toward the instantaneous spot on the mean spherical surface where specular reflection occurred (8 = 8i = 8r). The reflection process converts incident RCP waves to echoes with right circular polarization (RCP) and left circular polarization (LCP) components. A weak signal often reached Earth directly, radiated via an antenna sidelobe, with Doppler shift fd. Doppler shift along the specular point reflection path was fr. On Earth, the RCP and LCP signals were captured by separate receivers and digitally recorded for later processing. For Magellan bistatic radar, the significant difference was that the spacecraft radiated a linearly polarized wave. The baseline noise was subtracted from the total power to obtain the echo power. [11] Additionally, complex cross spectra between the RCP and LCP channels hVRVL*i were computed at both wavelengths (Figure 3). The magnitude of the cross spectrum is a measure of the fractional surface which behaves as specular reflectors, and its phase is sensitive to the argument of e. The cross spectra of the noise in the RCP and LCP channels showed about 1 percent correlation, presumably from imperfect polarization isolation between the ground antenna feeds. We used these measured correlations, calculated over intervals of 7 –15 min, to correct for the coupling between the feeds, thereby improving the effective feed isolation, our ability to measure polarization of the surface echoes, and our ability to infer the argument of e.
3. Analysis [12] Specular scattering from the surface converts an incoming monochromatic VEX RCP wave into an outgoing wave with both RCP and LCP components, the magnitudes and phases of which depend on the Fresnel reflection coefficients [Stratton, 1941]: RH ¼ cos 8 sqrt e sin2 8 = cos 8 þ sqrt e sin2 8 ð1aÞ RV ¼ e cos 8 sqrt e sin2 8 = e cos 8 þ sqrt e sin2 8 ð1bÞ
RH and RV are the reflection coefficients for the horizontal and vertical components of the incident electromagnetic wave and 8 is the angle of incidence. The RCP and LCP
echo components, when RCP is incident, are given by [Simpson, 1993] RR ¼ ðRV þ RH Þ=2 and RL ¼ ðRV RH Þ=2
ð2Þ
For partially conducting interfaces, equations (1) and (2) may be used by replacing the ‘‘real’’ dielectric constant with its complex equivalent, i.e., e — > e0 is/2pf = jej exp[i( f )], where e0 is the real part of the dielectric constant, s is the surface material conductivity, f is the wave frequency, and q( f ) is the argument of e. [13] The first step in finding the surface dielectric constant e from measurements is to solve (1) and (2) using the powers in the received RCP and LCP echo signals [Simpson et al., 2006]. The ratio of RCP to LCP power depends strongly on the magnitude of the dielectric constant but not its argument (Figure 4a). The angle at which RCP and LCP echo powers are equal, the Brewster angle for real e, increases from 60° to 75° to 84.3° for jej = 3, 10, and 100. [14] The phase of the received cross spectrum depends initially on the phase of the product of the circular reflection coefficients, RRRL*. With RCP incident on the surface, the difference in the two reflection coefficients imposes a phase shift between the RCP and LCP echo components. When the spacecraft rolls about the HGA axis during transmission, the phase difference between the RCP and LCP signals changes by 4p for each full rotation (VEX rolls for thermal protection). The two components may undergo further, small differential phase shifts as they pass through magnetized plasma between Venus and Earth. The separate paths through microwave circuits and electronics on the ground are certainly different as well. It is usually adequate to assume that spacecraft roll, plasma effects, and electronic drift add no more than a slowly time-varying bias to the
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Figure 2. Example VEX 13 cm RCP 1024-point 100 kHz power spectrum from 15 June. Spectrum is the average of 976 individual power spectra collected within ±5 s of 0156:16, when the specular point was near (78°N, 25°E) in a plains area northwest of Maxwell Montes. The surface echo is identified by the horizontal bar over 38– 47 kHz. Within the echo window a narrow carrier signal propagating directly from the spacecraft (Figure 1) is visible at 43 kHz. Baseline noise power density (No = 0.299 zeptowatts/Hz, where 1 zW is 1021 W) is calculated from the windows on either side of the echo. The carrier power, when present, is removed by linear interpolation across five frequency bins, after which the noise pedestal is removed and the echo power (Pe = 164.8 zW) is calculated by summing the remainders. The formal uncertainty in Pe (3.99 zW) is based on No and the amount of time and frequency averaging in the noise and echo windows. The uncertainty does not include the fluctuations in Pe, caused by the scattering process, which are typically 3 – 4 times larger (Table 3).
phase difference initiated by the RCP and LCP reflection coefficients. If the elevated regions on Venus are conducting while the surrounding plains are not, we would expect to see clear changes in the echo phase in hVRVL*i, mimicking the changes in the phase of hRRRL*i and occurring suddenly as the specular point moves into and out of the anomalously reflecting regions. [15] Figure 4b shows the predictions of cross-spectrum phase y from (1) and (2) for the nine choices of e (three each of magnitude and argument) used in Figure 4a. The cross-spectrum phase is ±180° for real dielectric constants. For imaginary e, y 135°, varying slightly with incidence angle 8 for the smallest e. The cross-spectrum phase is approximately the average of the real and imaginary values when e itself is midway between real and imaginary. Thus, to first order, y 180 q=2 þ bias terms
ð3Þ
The corresponding cases for Magellan’s 13 cm bistatic radar experiment are shown in Figure 5. Figure 5a shows the RCP/LCP power ratio for the three example choices of jej and three choices for q when the incident wave is linearly polarized at 45° with respect to the incidence plane. For the Magellan configuration the power ratio would be unity when the dielectric constant is real. If the dielectric constant were purely imaginary, RCP echo power can become 10 times larger than LCP near the Brewster angle for the corresponding real e, with the difference being most pronounced for the smallest jej. Figure 5b shows the phase of the echo cross spectrum, which is most sensitive to 8 and jej and least sensitive to q.
4. Data [16] Figure 6 shows example HGA ‘‘footprints’’ across Maxwell Montes for the bistatic radar experiments. The MGN footprints moved from southwest to northeast; since the experiments were conducted on alternating 93 min
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Figure 3. Example VEX 13 cm cross spectrum on 19 June at 0158:55 ± 05, when the specular point was near (76°N, 9°E). (a) The directly propagating carrier, visible at 43 kHz in Figure 2, can be seen in the magnitude plot at 73 kHz. (b) Phase values are clustered at y 100°, where the surface echo is well defined, but are distributed randomly elsewhere.
orbits, the tracks are nearly identical. The Venus Express footprints moved from northwest to southeast on alternating 24 h orbits. Of particular interest is the VEX track on 19 June, which crossed the MGN tracks near Cleopatra, where the specular echoes implying conducting surface material were obtained. Here we focus on 19 June. The results from 15 and 17 June and are similar but without the Cleopatra connection. [17] Figure 7 provides an overview of the 13 cm VEX results on 19 June. Details for Cleopatra Patera, Maxwell Montes, and the immediate surroundings are expanded in Figure 8 and listed in Table 3. The solid lines in Figure 7 show amplitudes and phases averaged over 25 frequency bins spanning 2441 Hz centered on the echo peak, consistent with the echo widths in Figures 2 and 3. Results for the peak bin alone (Df 98 Hz) are shown by the open circles. Averages of noise in 25 bins approximately 10 kHz above and below the surface echo are indicated by ‘‘plus’’ and ‘‘cross’’ symbols, respectively. [18] A predominantly specular echo is expected to have a cross-spectrum magnitude that is close to the geometric mean of the RCP and LCP echo powers and a crossspectrum phase that has a small standard deviation. As the scattering process becomes less specular, we expect the
correlation between RCP and LCP echo components to decrease, the cross-spectrum magnitude to drop below the RCP-LCP geometric mean, and the cross-spectrum phase to approach a uniform distribution. The extreme case of no echo is illustrated in the noise portions of spectra. For example, over 55 – 65 kHz in Figure 3, the RCP and LCP noise power densities are 0.299 and 0.416 zW/Hz, respectively, the cross-spectrum magnitude is small, and the phases are distributed randomly. In Figure 8a, the noise power has been subtracted to leave the RCP and LCP echo powers. The majority of measurements for t < 2.011 and t > 2.039 in Figure 8 satisfy the criteria for specular echoes in that the cross-spectrum magnitude is close to the RCP-LCP geometric mean and the crossspectrum phase is narrowly distributed. Also, measurements in this range have larger RCP than LCP echo power density, as is expected for reflection at incidence angles greater than the Brewster angle. [19] For 2.011 < t < 2.039, at least seven of the ten measurements in Figure 8a show LCP echo power density larger than RCP, implying a much larger e than in the surrounding terrain. Only the last measurement (t = 2.0375) meets the criteria for specular reflection, however. Half of the measurements have cross-spectrum magnitudes smaller
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Figure 4. (a) Ratio of hVRVR*i to hVLVL*i calculated from equations (1) and (2) for RCP incident (the VEX case), three choices of dielectric constant magnitude (jej = 3, 10, 100), and three choices of dielectric constant argument (q = 0°, 45°, and 90°). (b) Phase y of the cross spectrum hVRVL*i is given for the same conditions.
than both RCP and LCP echo powers, and only the last measurement has a cross-spectrum phase standard deviation less than 40° (Table 3, rows 5 – 14). Of particular concern is the cross spectrum that we associate most closely with Cleopatra (t = 2.0208; Table 3, row 8). There, a broad maximum in cross-spectrum magnitude (0.008 zW/Hz) is centered on Cleopatra; but the RCP and LCP echo power densities are also relatively large with values 0.018 and 0.025 zW/Hz, respectively (Figure 8a). Taken together, these measurements imply that less than 40% of the echo power is from specular scattering. The standard deviation of the cross-spectrum phase at t = 2.0208 is 40°, the second smallest within the Maxwell group, and not unreasonable if there is a mix of specular and ‘randomizing’ elements within the illuminated footprint. The cross-spectrum phase itself is 92°, midway between the four-row averages at the beginning (106°) and end (82°) of Table 3. Although 92° could be taken as evidence against a sudden shift in electrical properties, we note that the measurements immediately before and after yielded cross-spectrum phases of 154 ± 51° (Table 3, row 7) and 127 ± 44° (row 9), which diverge from the linear trend apparent in Figure 7d. However, all three values have overlapping one-sigma uncertainties. [20] To calculate the total echo power Pe and its uncertainty in each polarization we used the method illustrated in
Figure 2. The dielectric constant can then be obtained by solving (1) and (2). In practice, we used the numerical RCP/ LCP ratio and the incidence angle 8 to obtain e by interpolation from a precomputed lookup table. A total of nine solutions for e was sought for each (RCP, LCP) pair of Pe values (Table 3, Notes). When any of the nine (RCP, LCP) pairs failed to yield a solution, ‘N/A’ was entered in the table. [21] Before solving for e, we adjusted the incidence angle for the effects of atmospheric refraction, including differences in refraction resulting from surface elevation changes (specular point latitude and longitude are changed very little by refraction). The incidence angle correction is as much as 4° ten minutes before the specular point entered Maxwell Montes. The maximum topography of Maxwell reduces a 2.3° correction to 1.3°. The observational geometry is almost identical among the three VEX orbits except for a 30 s/day delay in time and a 1.8°/day eastward shift in longitude of the specular point, resulting primarily from the rotation of Venus. [22] The 10 s spacing of data points in Table 3 represents a compromise between long integration times needed to improve signal-to-noise ratio and short integration times needed to preserve temporal and spatial resolution. Averages over 30 s consistently resulted in solutions for e, but the enhanced echoes from Cleopatra were not well resolved.
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Figure 5. (a) Ratio of hVRVR*i to hVLVL*i when a linearly polarized wave is incident for three choices of dielectric constant magnitude (jej = 3, 10, 100) and three choices of the dielectric constant argument (q = 0°, 45°, and 90°). The incident angle of polarization is halfway between horizontal and vertical: the Magellan case. Note all ratios approach unity for small angles of incidence. In order to facilitate display, we have displaced the curves for jej = 3 upward and the curves for jej = 100 downward. (b) Phase y of the cross spectrum hVRVL*i under the same conditions, but without displacements.
The 10 s points are separated by about 85 km along the specular track (Figure 6). The circularly polarized antenna beam from VEX was about 84 km across, half power to half power, where it intercepted the surface at periapsis (slant range 820 km on 19 June at t = 2.033). The HGA pattern, projected onto the surface at 8 = 71°, results in an elliptical footprint with dimensions approximately 84 260 km; five minutes earlier and seven minutes later the ellipse dimensions were doubled, chiefly because of increased slant range. The long axis of the ellipse was approximately perpendicular to the specular point track during the experiment, so the measurements listed in Table 3 can be considered independent.
5. Discussion 5.1. Magellan [23] We reprocessed the Magellan bistatic radar data [Simpson, 1997] using the same software that was applied to the VEX BSR data. One source of concern has been a possible ground receiver problem, which might have resulted in an apparent excess in Magellan RCP echo power and led to the conclusion that the surface near Cleopatra is
electrically conducting. We also know that the ground antenna system is imperfect. However, our experience since 2004 with a number of bistatic radar experiments and tests at DSN 70 m antennas is that cross polarization between RCP and LCP is approximately 25 dB below the direct polarization level. [24] Our reprocessing of the Magellan data included calibration of the absolute power levels. Compared with an original ad hoc assumption that the two receiver channels had equal gains and system noise (neither of which was known), the reprocessed results are less stable than the originals because of variations in system gain (especially LCP) of up to 1 dB on time scales of seconds to minutes. Also, because RCP and LCP are not balanced, the differences between channels appear as a weakly polarized (and variable) component of the noise where none should be expected. These instabilities dominate the physical cross polarization in the Magellan data in both an average sense and in terms of minute-to-minute fluctuations. [25] When the specular point is outside Maxwell, typical Magellan echoes have power densities larger than 0.1 zW/Hz (Figure 9a). Because the transmitted polarization was linear, both RCP and LCP have similar echo powers at
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Figure 6. Magellan and Venus Express ground tracks, shown as a sequence of antenna beam footprints superimposed on a topographic map of the Maxwell area [U.S. Geological Survey, 1984]. Footprints are separated by 30 s. Magellan footprints moved from southwest to northeast along three almost identical tracks spaced by 186 min in time (only one orbit is shown). Venus Express footprints moved from northwest to southeast on 15 –19 June (only the first and last orbits are shown).
the receiver; the power ratio is unity for latitudes less than 60° and larger than 70° (Figure 9b). The echo powers from within Maxwell drop to about a tenth those from areas immediately adjacent to Maxwell. Nonetheless, it is possible to identify the excess RCP power during the latter half of the Maxwell passage (latitudes 65– 67° in Figure 9), which Pettengill et al. [1996] associated with Cleopatra. Although this excess may sometimes be associated with the short-term gain fluctuations, the amount of excess is clearly different in spectral regions occupied by echo signals as compared with nearby noise. In addition, the RCP excess recurs on each of the three Magellan orbits over Cleopatra Patera, and no similar feature appears in echoes at other surface locations. On this basis, we support the original conclusion by Pettengill et al. [1996] regarding the excess RCP power in the Maxwell echoes. [26] From Figure 5 we can confirm that the 100° crossspectrum phase change over Maxwell (Figure 9c) is consistent with jej 100 (or larger); a power ratio of 1.4 is consistent with equal real and imaginary parts, while a power ratio of 1.8 favors an entirely imaginary dielectric constant. Neither of these results is strongly dependent on incidence angle, which is approximately 2° larger than the latitude during this time.
5.2. Applicability of Fresnel Reflection [27] The methods presented here are based on an assumption that the surface echo from Venus is dominated by Fresnel reflection, equations (1) and (2), reflection by many surface ‘facets’ having radii of curvature that are large compared to the wavelength [Tyler, 1976]. Over most of Venus, the magnitude of the cross spectrum is near the geometric mean of the RCP and LCP powers and the crossspectrum phase is well defined, validating the assumption. The data from Cleopatra Patera, however, imply that less than 40% of the echo originates from such surface elements. Around Cleopatra the model may be inadequate, and over other parts of Maxwell the breakdown may be even more severe. [28] However, the VEX data do show that there is a reversal of the sense of circular polarization over Maxwell generally and near Cleopatra specifically. We do measure a nonzero cross spectrum with constrained (albeit not small) phase distributions: at least in the vicinity of Cleopatra. These are results we would expect from Fresnel reflection even though we admit that (1) and (2) cannot be applied rigorously. A literature of Fresnel alternatives exists, scattering by discrete objects (e.g., wires and rocks) and
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Figure 7. Summary results from Venus Express 13 cm bistatic radar experiment on 19 June. (a) RCP power density, (b) LCP power density, (c) cross-spectrum magnitude, and (d) and cross-spectrum phase. Horizontal axis is Earth receive time in hours past midnight. Solid lines are average values over 25 frequency bins (2441 Hz) centered on the surface echo peak. Open circles are values in the central echo bin (Df = 98 Hz). Plus and cross symbols are 25 bin averages 10 kHz above and below the echo, respectively, where only noise is present. Note in Figure 7d that surface echo phase drifts toward smaller values during the experiment because of spacecraft attitude and other slow changes. For more detail on Maxwell Montes, including Cleopatra Patera, see Figure 8 and Table 3.
irregular structure (edges and cracks), but the results do not explain the Venus observations. Dedrick et al. [1978] have shown that the cross-polarized radar cross section of an ‘‘atmosphere’’ of randomly oriented wires is one-third that of the direct cross section in both the forward and backward directions. This sets a limit on the amount of polarization change that can be attributed to discrete objects and fails to explain the polarization reversal we see at Maxwell Montes. Refraction has been proposed by Tryka and Muhleman [1992] and Peters [1992] to explain unexpected phenomena on Venus and icy satellites, respectively, but the models require a regolith which is nearly lossless and are more attractive for backscatter. [29] Cost [1965] and Ulaby et al. [1988] have studied the microwave bistatic response of terrestrial surfaces both with
and without vegetation. Their experiments were typically conducted in a controlled (outdoor laboratory) environment with orthogonal linear polarizations. This and other work has been summarized by Weiner [2007]. Almost without exception the maximum bistatic radar cross section of these surfaces was observed to be consistent with specular theory: even for ‘rough’ surfaces. For incidence angles 8 > 40° the response in the specular direction was typically 10 dB or more than in other directions. [30] Lacking a detailed, viable alternative model we propose that many of the key features of Fresnel reflection are being reproduced in Maxwell Montes by its very rough surface. Although (1) and (2) cannot be applied because none of the ‘facets’ is large enough to qualify as a quasi-
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Figure 8. The 13 cm VEX bistatic radar echo measurements collected near and within Maxwell Montes. Maxwell is assumed to extend over 2.011 < t < 2.039, with Cleopatra Patera at t 2.02. (a) RCP (plusses) and LCP (crosses) echo power density averaged over 25 frequency bins centered on the echo peak; noise power has been subtracted from total power listed in Table 3, columns 6– 7. Cross-spectrum magnitude density (circles) is taken from Table 3, column 8. (b) Cross-spectrum average phase (circles) and its standard deviation from calculations over the same 25 frequency bins.
specular reflector, the bulk properties of the surface material may still be sensed. A surface saturated with relatively flat rocks would have a hybrid scattering response, partly from the dipole nature of the rocks and partly from the exposed flat side, which can be responsible for polarization reversal in the bistatic geometry. Working the problem backward, we can interpret changes in the RCP/LCP ratio in terms of the material dielectric constant without being able to say exactly what the values are. We have previously investigated the scattering by populations of discrete rocks on and within a regolith [Baron et al., 1998]. Similar methods may be applied here; but the model must incorporate rock-rock interactions, which makes an already computationally intensive calculation an order of magnitude more difficult.
6. Conclusions [31] The three Venus Express bistatic radar experiments conducted over Maxwell Montes confirm the presence of high dielectric material in elevated regions through measurements of opposite sense, circularly polarized echo powers when a circularly polarized wave was incident.
The echo power ratio yielded values in the range 4 < jej < 25 for Maxwell, but the weak echoes and relatively low cross-spectrum magnitude caution against a strict Fresnel-based interpretation. Away from Maxwell Montes, where criteria for specular scattering are more easily met, values jej 4 are common. [32] Attempts to detect electrical phase changes in the VEX cross spectrum hVRVL*i, expected if the surface material were conducting, have not been successful. Magellan observations of excess RCP power, which require that e have a large imaginary component, have been confirmed near Cleopatra crater. VEX results from the same area are inconclusive because of inadequate signal-to-noise ratio. [33] Continuing the VEX BSR explorations, particularly over areas other than Maxwell and at smaller Earth-Venus distances remains a desirable goal. The VEX opportunity has been lost because of failure in the spacecraft radio system, which after November 2006 radiated only 200 mW at 13 cm wavelength. Future electromagnetic probing of the unusual high-elevation terrains on Venus must await the arrival of a new spacecraft.
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2.0014 2.0042 2.0069 2.0097 2.0125 2.0153 2.0181 2.0208 2.0236 2.0264 2.0292 2.0319 2.0347 2.0375 2.0403 2.0431 2.0458 2.0486
1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18
71.52 71.10 70.95 70.73 70.46 70.17 69.87 69.55 69.15 68.77 68.42 67.93 67.40 66.83 66.32 65.88 65.51 65.07
8 (°) 71.73 71.00 70.27 69.52 68.76 67.98 67.20 66.41 65.60 64.79 63.98 63.15 62.31 61.47 60.63 59.79 58.96 58.13
Lat (°N) 2.29 3.38 4.37 5.28 6.12 6.83 7.50 8.11 8.63 9.11 9.56 9.92 10.25 10.54 10.80 11.06 11.30 11.50
Lon (°E) 0.334 0.334 0.313 0.306 0.305 0.306 0.307 0.319 0.312 0.304 0.303 0.303 0.306 0.312 0.356 0.325 0.308 0.313
± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±
0.016 0.018 0.009 0.010 0.011 0.010 0.009 0.010 0.007 0.012 0.009 0.013 0.010 0.011 0.030 0.012 0.010 0.014
RCPd 0.441 0.434 0.417 0.418 0.429 0.424 0.431 0.443 0.434 0.425 0.419 0.421 0.430 0.434 0.464 0.438 0.429 0.426
± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±
0.021 0.017 0.011 0.011 0.010 0.010 0.013 0.012 0.015 0.010 0.016 0.014 0.013 0.015 0.025 0.013 0.014 0.015
LCPd 0.029 0.027 0.009 0.007 0.003 0.004 0.007 0.008 0.008 0.006 0.002 0.001 0.001 0.015 0.052 0.022 0.015 0.015
± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±
0.012 0.012 0.007 0.006 0.009 0.009 0.009 0.012 0.008 0.007 0.008 0.008 0.006 0.011 0.026 0.011 0.006 0.011
Peak Magnitude (zW/Hz) 110.2 ± 17.0 110.0 ± 18.3 93.4 ± 46.6 111.7 ± 50.2 74.6 ± 69.8 114.9 ± 60.6 154.5 ± 51.3 92.2 ± 40.0 126.5 ± 43.6 101.7 ± 49.2 137.7 ± 79.7 132.9 ± 85.7 25.9 ± 79.1 78.2 ± 21.9 83.3 ± 7.4 79.4 ± 15.0 86.1 ± 29.0 79.8 ± 26.2
Phase (°)
Cross-Spectrum Average
131.23 ± 10.87 96.69 ± 10.95 24.64 ± 10.89 62.69 ± 10.83 4.23 ± 10.86 43.22 ± 10.77 23.07 ± 10.83 60.97 ± 10.79 52.10 ± 10.73 36.38 ± 10.75 6.48 ± 10.75 12.53 ± 10.72 36.31 ± 10.68 54.49 ± 10.66 162.26 ± 10.64 100.07 ± 10.56 54.23 ± 10.54 72.33 ± 10.54
RCPt
85.28 ± 15.28 70.63 ± 15.16 36.92 ± 15.18 15.46 ± 15.11 75.36 ± 15.10 54.95 ± 15.03 37.68 ± 15.12 130.38 ± 14.98 70.14 ± 15.06 69.00 ± 14.97 41.05 ± 14.94 45.06 ± 14.83 40.92 ± 14.86 103.63 ± 14.77 187.01 ± 14.73 94.71 ± 14.74 39.56 ± 14.73 47.43 ± 14.73
LCPt
Total Echo Power per Figure 2 (zW)
e 6.150 ± 0.769 6.511 ± 1.014 N/A N/A N/A 9.849 ± 2.529 13.252 ± 6.386 14.619 ± 2.185 9.172 ± 1.865 12.363 ± 3.274 N/A 47.770 ± 51.2 6.682 ± 2.098 9.809 ± 1.640 5.877 ± 0.383 4.791 ± 0.554 3.803 ± 0.938 3.347 ± 0.646
From RCPt/LCPt
Columns 1 – 2: Values in each row are 10 s averages centered on the time in column 2. Columns 3 – 5: Geometry was corrected for atmospheric refraction (including differential refraction effects of topography) using a ray tracing model, the Magellan model atmosphere [Lyons, 1991], and topography derived from Magellan radar altimetry [Ford, 1992]. Columns 6 – 7: Power density and its standard deviation were computed using 25 frequency bins (2441 Hz) centered on the fitted position of the echo peak; of these totals, 0.299 – 0.301 and 0.416 – 0.420 zW/Hz were RCP and LCP noise power, respectively (see Figure 2). Columns 8 – 9: Cross-spectrum magnitude and phase averaged over 25 bins. Statistics were computed using real and imaginary parts; surface echoes are associated with small phase standard deviations, while noise has phase standard deviations 90°. Columns 10 – 11: Echo powers as calculated from components illustrated in Figure 2. Column 12: Dielectric constant was derived by solving equations (1) and (2) (see text) for each of nine RCP/LCP ratios and the corrected incidence angle (column 3). The nine ratios were constructed by pairing RCP power, RCP power less one standard deviation, and RCP power plus one standard deviation with the corresponding LCP values. Solutions were given a weighting of e0.5 for each occurrence of a standard deviation in the ratio. Values in column 12 are the weighted mean and standard deviation from the nine solutions. ‘N/A’ appears if any of the nine solution attempts failed. Rows 5 – 14 are associated with Maxwell Montes, row 8 with Cleopatra Patera.
a
Time (Hours Past 0 h)
Row
Average Power Density Noise + Echo Peak (zW/Hz)
Table 3. Venus Express Bistatic Radar Power and Cross-Spectrum Results 13 cm Wavelength 19 June 2006a
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Figure 9. Magellan bistatic radar results after reprocessing in the Maxwell Montes region, plotted versus specular point latitude, which is about 2° less than angle of incidence. Echo properties were averaged over 25 frequency bins (610 Hz at sampling rate 25 kHz). (a) Echo power from the sum of the RCP and LCP power densities. (b) Ratio of RCP to LCP power densities, expected to be near 1.0 for surfaces with real dielectric constants. (c) Cross-spectrum phase. Cleopatra Patera (65 – 67° latitude) is unique in having moderately strong echoes, excess RCP power (ratio 1.5), and a large phase change.
[34] Acknowledgments. This work was supported at Stanford University by the NASA Venus Express Participating Scientist Program (grant NNG06GG04G-000001). Estimates of absorption by the Venus atmosphere were based on a model kindly provided by J.M. Jenkins of the SETI Institute. Comments on an earlier version of the manuscript by Peter Ford, Gordon Pettengill, and an anonymous reviewer are gratefully acknowledged.
References Baron, J. E., R. A. Simpson, G. L. Tyler, H. J. Moore, and J. K. Harmon (1998), Estimation of Mars radar backscatter from measured surface rock populations, J. Geophys. Res., 103, 22,695 – 22,712, doi:10.1029/ 98JE02221. Cost, S. T. (1965), Measurements of the bistatic echo area of terrain at X-band, Rep. 1822 – 2, Ohio State Univ. Antenna Lab., Columbus. Dedrick, K. G., A. R. Hessing, and G. L. Johnson (1978), Bistatic radar scattering by randomly oriented wires, IEEE Trans Antennas Propag., 26, 420 – 426. Ford, P. G. (1992), MGN V RDRS 5 Global Data Record Topographic V1.0 (MGN-V-RDRS-5-GDR-TOPOGRAPHIC-V1.0), NASA Planet. Data Syst., Greenbelt, Md. Ford, P. G., and G. H. Pettengill (1983), Venus: Global surface radio emissivity, Science, 220, 1379 – 1381, doi:10.1126/science.220.4604.1379. Goldstein, R. M. (1965), Preliminary Venus radar results, J. Res. Natl. Bur. Stand., Sect. D, 69, 1623 – 1625. Ha¨usler, B., et al. (2006), Radio science investigations by VeRa onboard the Venus Express spacecraft, Planet. Space Sci., 54, 1315 – 1335, doi:10.1016/j.pss.2006.04.032. Jurgens, R. F. (1970), Some preliminary results of the 70-cm radar studies of Venus, Radio Sci., 5, 435 – 442, doi:10.1029/RS005i002p00435. Klose, K. B., J. A. Wood, and A. Hashimoto (1992), Mineral equilibria and the high radar reflectivity of Venus mountaintops, J. Geophys. Res., 97, 16,353 – 16,369, doi:10.1029/92JE01865.
Lyons, D. T. (1991), Magellan planetary constants and models, JPL D-2300 Rev. D, Jet Propul. Lab., Pasadena, Calif. Peters, K. J. (1992), The coherent backscatter effect: A vector formulation accounting for polarization and absorption effects and small or large scatterers, Phys. Rev. B, 46, 801 – 812, doi:10.1103/PhysRevB.46.801. Pettengill, G. H. (1962), Enhancement of radar reflectivity associated with the lunar crater Tycho, J. Geophys. Res., 67, 4881 – 4885, doi:10.1029/ JZ067i012p04881. Pettengill, G. H., E. Eliason, P. G. Ford, G. B. Loriot, H. Masursky, and G. E. McGill (1980), Pioneer Venus radar results: Altimetry and surface properties, J. Geophys. Res., 85, 8261 – 8270, doi:10.1029/ JA085iA13p08261. Pettengill, G. H., P. G. Ford, and R. J. Wilt (1992), Venus surface radiothermal emission observed by Magellan, J. Geophys. Res., 97, 13,091 – 13,102, doi:10.1029/92JE01356. Pettengill, G. H., P. G. Ford, and R. A. Simpson (1996), Electrical properties of the Venus surface from bistatic radar observations, Science, 272, 1628 – 1631, doi:10.1126/science.272.5268.1628. Rogers, A. E. E., and R. P. Ingalls (1970), Radar mapping of Venus with interferometric resolution of the range-Doppler ambiguity, Radio Sci., 5, 425 – 433, doi:10.1029/RS005i002p00425. Schaefer, L., and B. Fegley Jr. (2004), Heavy metal frost on Venus, Icarus, 168, 215 – 219, doi:10.1016/j.icarus.2003.11.023. Shepard, M. K., R. E. Arvidson, R. A. Brackett, and B. Fegley Jr. (1994), A ferroelectric model for the low emissivity highlands on Venus, Geophys. Res. Lett., 21, 469 – 472, doi:10.1029/94GL00392. Simpson, R. A. (1993), Spacecraft studies of planetary surfaces using bistatic radar, IEEE Trans. Geosci. Remote Sens., 31, 465 – 482. Simpson, R. A. (1997), Magellan Bistatic Radar Raw Data Records V1.0 (MGN-V-RSS-1-BSR-V1.0), NASA Planet. Data Syst., Greenbelt, Md. Simpson, R. A., G. L. Tyler, M. Pa¨tzold, and B. Ha¨usler (2006), Determination of local surface properties using Mars Express bistatic radar, J. Geophys. Res., 111, E06S05, doi:10.1029/2005JE002580. Stratton, J. A. (1941), Electromagnetic Theory, McGraw Hill, New York.
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Tryka, K. A., and D. O. Muhleman (1992), Reflection and emission properties on Venus: Alpha Regio, J. Geophys. Res., 97, 13,379 – 13,394, doi:10.1029/92JE01163. Tyler, G. L. (1976), Wavelength dependence in radio-wave scattering and specular-point theory, Radio Sci., 11, 83 – 91, doi:10.1029/ RS011i002p00083. Tyler, G. L., and D. H. H. Ingalls (1971), Functional dependencies of bistatic-radar frequency spectra and cross sections on surface scattering laws, J. Geophys. Res., 76, 4775 – 4785, doi:10.1029/JB076i020p04775. Ulaby, F. T., T. E. Van Deventer, J. R. East, T. F. Haddock, and M. E. Coluzzi (1988), Millimeter-wave bistatic scattering from ground and vegetation targets, IEEE Trans. Geosci. Remote Sens., 26, 229 – 243. U.S. Geological Survey (1984), Topographic and shaded relief maps of Venus (map I-1562), Reston, Va.
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Weiner, M. W. (2007), Clutter, in Advances in Bistatic Radar, edited by N. J. Willis and H. D. Griffiths, chap. 9, pp. 230 – 319, SciTech, Raleigh.
B. Ha¨usler and R. Mattei, Institut fu¨r Raumfahrttechnik, Universita¨t der Bundeswehr Mu¨nchen, D-85577 Neubiberg, Germany. (bernd.haeusler@ unibw-muenchen.de;
[email protected]) G. L. Tyler and R. A. Simpson, Space, Telecommunications, and Radioscience Laboratory, Stanford University, Packard Building, MC 9515, Stanford, CA 94305-4020, USA. (
[email protected]; rsimpson@ magellan.stanford.edu) M. Pa¨tzold, Rheinisches Institut fu¨r Umweltforschung, Abt. Planetenforschung, Aachener Strasse 209, D-50931 Ko¨ln, Germany. (mpaetzol@ uni-koeln.de)
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, E00B24, doi:10.1029/2008JE003134, 2008
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Felsic highland crust on Venus suggested by Galileo Near-Infrared Mapping Spectrometer data George L. Hashimoto,1 Maarten Roos-Serote,2 Seiji Sugita,3 Martha S. Gilmore,4 Lucas W. Kamp,5 Robert W. Carlson,5 and Kevin H. Baines5 Received 2 March 2008; revised 29 July 2008; accepted 18 September 2008; published 31 December 2008.
[1] We evaluated the spatial variation of Venusian surface emissivity at 1.18 mm
wavelength and that of near-surface atmospheric temperature using multispectral images obtained by the Near-Infrared Mapping Spectrometer (NIMS) on board the Galileo spacecraft. The Galileo NIMS observed the nightside thermal emission from the surface and the deep atmosphere of Venus, which is attenuated by scattering from the overlying clouds. To analyze the NIMS data, we used a radiative transfer model based on the adding method. Although there is still an uncertainty in the results owing to the not well known parameters of the atmosphere, our analysis revealed that the horizontal temperature variation in the near-surface atmosphere is no more than ±2 K on the Venusian nightside and also suggests that the majority of lowlands likely has higher emissivity compared to the majority of highlands. One interpretation for the latter result is that highland materials are generally composed of felsic rocks. Since formation of a large body of granitic magmas requires water, the presence of granitic terrains would imply that Venus may have had an ocean and a mechanism to recycle water into the mantle in the past. Citation: Hashimoto, G. L., M. Roos-Serote, S. Sugita, M. S. Gilmore, L. W. Kamp, R. W. Carlson, and K. H. Baines (2008), Felsic highland crust on Venus suggested by Galileo Near-Infrared Mapping Spectrometer data, J. Geophys. Res., 113, E00B24, doi:10.1029/2008JE003134.
1. Introduction [2] Venus is the nearest neighbor of Earth, and these twins are close in mass, size, and bulk composition. However, the surface environment of Venus is completely different from that of Earth, where the mean surface temperature is 735 K and water is almost absent. In contrast, the Earth surface is covered with ocean and is habitable. It is not clear whether the surface environments on these twin planets differed from beginning or they evolved from a similar environment into the current different states. Understanding the evolution of these twin planets is one of the most important and elusive problems in planetary climatology. [3] The chemical and mineralogical composition of the Venus surface can provide valuable information on the climatic evolution of Venus. If geologic signatures that are produced only under unique conditions are found in the chemical and mineralogical composition, they could be used 1 Laboratory for Earth and Planetary Atmospheric Science, Organization of Advanced Science and Technology, Kobe University, Kobe, Japan. 2 Lisbon Astronomical Observatory, Lisbon, Portugal. 3 Department of Complexity Science and Engineering, Graduate School of Frontier Science, University of Tokyo, Kashiwa, Japan. 4 Department of Earth and Environmental Sciences, Wesleyan University, Middletown, Connecticut, USA. 5 Jet Propulsion Laboratory, California Institute of Technology, Pasadena, California, USA.
Copyright 2008 by the American Geophysical Union. 0148-0227/08/2008JE003134$09.00
to reconstruct the Venus paleoenvironment. Additionally, a global map of surface composition will be a very important diagnostic of bulk planetary composition, chemical differentiation, and evolution of the planetary interior. [4] It has been also argued that the controlling mechanism of the present Venus surface environment is a connection between the atmosphere and the chemical characteristics of planetary surface [e.g., Hashimoto and Abe, 2005, and references therein]. The high surface temperature of Venus enhances the importance of the gas-solid reactions on the planetary surface in controlling the atmospheric composition, including infrared-active species which control the greenhouse effect. Venus climate models that combine chemical reaction and radiative transfer in the atmosphere have demonstrated that such a coupling between surface chemical reactions and greenhouse effect will play an important role in controlling the surface environment [e.g., Hashimoto and Abe, 2000; Bullock and Grinspoon, 2001]. [5] However, chemical characteristics of the Venus surface are not well constrained. In situ measurements of the surface using gamma-ray and X-ray fluorescence spectroscopy were performed by Venera and Vega landers [Surkov et al., 1984, 1986, 1987], but it is difficult to infer the global characteristics from lander data collected at 7 sites on the planet. Although the radar reflectivity and radiothermal emissivity were measured nearly globally [e.g., Pettengill et al., 1982, 1992], origins of the microwave emissivity differences are highly controversial and appear to be related to both grain size and dielectric constant of constituent
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minerals [e.g., Campbell et al., 1997; Pettengill et al., 1997]. We need a new observation capable of determining the chemical and mineralogical composition of the Venus surface. [6] It has been discovered that thermal emission from the surface of Venus is detectable at the CO2-free windows in the near-infrared wavelengths (0.85, 0.90, 1.01, 1.10, and 1.18 mm) [e.g., Carlson et al., 1993a, 1993b; Meadows and Crisp, 1996; Baines et al., 2000]. Thermal emissivity is a function of surface temperature, grain size, and mineral composition. Previous investigations of surface emissivity using Earth-based telescopic data have demonstrated variations that appear to be correlated with topography and therefore surface temperature [Lecacheux et al., 1993; Meadows and Crisp, 1996]. Although the variations in the surface emissivity have been investigated, no signature of surface emissivity differences was found at near-infrared wavelengths in contrast with the microwave wavelengths [Lecacheux et al., 1993; Meadows and Crisp, 1996]. However, these analyses neglect multiple reflection between the planetary surface and the clouds which has been demonstrated to significantly obscure variations in surface emissivity [Moroz, 2002; Hashimoto and Sugita, 2003]. Reexamination of previous analyses demonstrated that there may be a significant spatial variation in the surface emissivity as large as 20% [Hashimoto and Sugita, 2003]. [7] Here in this study, we evaluated the Venus surface emissivity at 1.18 mm wavelength using data obtained by the Near-Infrared Mapping Spectrometer (NIMS) aboard the Galileo spacecraft. Because emissivity depends on the mineralogy, a mapping of surface emissivity may constrain the chemical characteristics of the Venus surface. In the following, we first describe the procedure of our analysis which includes the reflection by the planetary surface. Then, we evaluate the surface emissivity, and discuss the errors and assumptions that may affect the estimation. Finally, the existence of paleo-ocean on early Venus is inferred from the spatial variation of surface emissivity.
2. Data and Analysis 2.1. Data [8] We analyzed the multispectral nightside image identified as VPDIN-1 (for Venus Partial Disk Imaging-NightSide) which covers the disk from approximately 20°W to 90°E longitude [Carlson and Taylor, 1993] and is distributed by the Planetary Data System (PDS). Those data were obtained by Galileo NIMS during its flyby of Venus on 10 February 1990. NIMS is an imaging spectrometer, operating in the spectral range from 0.7 to 5.2 mm [Carlson et al., 1992], and images of nightside are obtained with 17 spectral channels [Carlson et al., 1991]. Some channels measured the thermal radiation in the well-known spectral windows at 1.18, 1.74, and 2.3 mm. In these windows, thermal emission originating from both the surface and lower atmosphere are observed [e.g., Allen and Crawford, 1984; Crisp et al., 1989; Carlson et al., 1991; Lecacheux et al., 1993; Pollack et al., 1993; Baines et al., 2000]. [9] NIMS channel 3 observed the radiation of the 1.18 mm window, in which thermal emission from the surface leaks through the Venus clouds (Figure 1a). Initial results from the VPDIN observations are described by Carlson et al.
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[1991], which mention that the high-altitude (12 km above mean planetary radius) surface feature Maxwell Montes is detectable in the NIMS channel-3 image. The dark features near the top and the right of the image are likely due to the Ishtar and Aphrodite terra, respectively (Figures 1a and 1e). However, there are also some features irrelevant to the surface. They are probably produced by the effects of scattering by the Venus clouds, which is the dominant effect in channels 5 and 7 (Figures 1c and 1d). Other effects detectable in channel 3 are scattered sunlight and limb darkening. These effects need to be corrected separately. 2.2. Overview of Analysis [10] The analysis procedure for evaluating surface emissivity is the following. (1) We first remove the contribution of scattered sunlight, (2) we make a correction for emission angle, and (3) we evaluate the surface emissivity. Since the radiation at 1.18 mm window is affected by both clouds and surface topography, these contributions are evaluated to isolate the contribution of surface composition to thermal emissivity. [11] In this analysis procedure, we use the NIMS channels 4, 5, and 7 in addition to channel 3. Figure 2 provides a schematic representation of the characteristics of radiation observed by the four channels. NIMS channel 4 is used to remove the contribution of scattered sunlight, and the NIMS channels 5 and 7 are used to estimate the cloud properties. 2.3. Radiative Transfer Model [12] We calculated synthetic spectra by means of a lineby-line radiative transfer program including both scattering and absorption. The computational code was developed by G. L. Hashimoto and others at the University of Tokyo. The algorithm known as the adding method is used to calculate the radiation in a vertically inhomogeneous plane-parallel atmosphere [Goody and Yung, 1989]. The number of radiative streams is 12 in each hemisphere. Taking into account a triangular spectral bandpass with the width of 0.025 mm [Carlson et al., 1992], we calculate radiances for each channel from synthesized spectra in which molecular absorption features are resolved. [13] Following the method by Pollack et al. [1993], opacities of gases and particles were calculated. In all our simulations, we used the vertical profiles of gas mixing ratios developed by Pollack et al. [1993]. Line absorption data for CO2, H2O, CO, SO2, HF, and OCS were taken from the HITEMP and HITRAN2000 [Rothman et al., 2003; L. S. Rothman et al., HITEMP, the high-temperature molecular spectroscopic database, manuscript in preparation, 2008]. Total internal partition sums (TIPS) are calculated by the TIPS code developed by Fischer et al. [2003]. Rayleigh scattering cross sections for atmospheric molecules are calculated from the refractive indices [Keady and Kilcrease, 2000; Penndorf, 1957; Vardavas and Carver, 1984]. Optical properties of cloud particles are calculated by the Mie scattering code [Bohren and Huffman, 1983] based on the baseline cloud model [Pollack et al., 1993] and the optical constants of 75% H2SO4 solution [Palmer and Williams, 1975]. [14] We also introduced a continuum absorption to reproduce an acceptable fit to the NIMS data, since an additional source of continuum opacity is present owing to the far
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Figure 1. Multispectral images of nightside of Venus. These data were obtained during the first NIMS imaging sequence (VPDIN-1) at a resolution of 50 km/pixel. (a) An image at 1.18 mm (channel 3). (b) An image at 1.47 mm (channel 4). (c) An image at 1.74 mm (channel 5). (d) An image at 2.31 mm (channel 7). (e) A Magellan topographic map. This image covers from 20°W to 90°E of longitude. The topographic high near the north pole is the Ishtar Terra. The western edge of Ovda Regio can be seen near the rightmost part of the image. wings of strong CO2 bands and collision-induced CO2 opacity [Pollack et al., 1993]. Binary absorption coefficients for continuum opacity are determined to match the measured flux with the cloud model of Pollack et al. [1993]. The values
used in our simulations are 1.0 109 cm1 amagat2 for 1.18 mm window, 1.0 108 cm1 amagat2 for 1.74 mm window, and 7.0 108 cm1 amagat2 for 2.3 mm window, respectively, though these values are somewhat different
Figure 2. A schematic picture of radiation in Venus’s atmosphere and clouds. 3 of 10
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from those used in other analyses [e.g., Be´zard et al., 1990; Pollack et al., 1993]. 2.4. Scattered Sunlight [15] Nightside emissions at the near-infrared wavelengths contain scattered sunlight reflected off the dayside of the planetary disk. In particular, a strong component that brightens gradually toward the terminator is found in channel 4 (1.47 mm) (Figure 1b). Atmospheric radiative transfer models demonstrate that Venus does not emit thermal radiation at this wavelength. The Venusian CO2 atmosphere is too opaque to allow radiation to penetrate the atmosphere, though the surface and lower atmosphere of Venus are hot enough to emit thermal radiation at this wavelength. The radiance observed in channel 4 can serve as a good approximation for the scattered sunlight. [16] We corrected the effect of scattered sunlight, by assuming that the contributions of this effect to channel 3, 5, and 7 are proportional to that of channel 4. Taking into account the solar spectrum, we use the following correction: IxS ¼ Ix
Sx I4 S4
ð1Þ
where suffix denotes the NIMS channel, ISx is the radiance after the correction of scattered sunlight, Ix is the radiance observed by the NIMS, and Sx is the solar spectrum. 2.5. Limb Darkening [17] We computed limb darkening by using our radiative transfer model. Figure 3 shows our calculated limb darkening curve and NIMS observation that is corrected for the scattered sunlight. The observed limb darkening is well reproduced by our simulation. Although the particle size and number density in the lower clouds would vary in time and space, they affect only scales of radiance and do not affect the shape of limb darkening. Since the upper cloud layer is opaque enough to determine the distribution of emission angle at the cloud top, the shape of limb darkening curve is independent of the property of the lower clouds. [18] The equation for limb darkening correction is given by IxSL ¼
Jx ð0Þ S Ix Jx ðqÞ
ð2Þ
where ISL x is the radiance after the correction of scattered sunlight and limb darkening, Jx (q) is the computed radiance for emission angle of q, respectively. 2.6. Estimation of Surface Emissivity [19] To evaluate the surface emissivity at 1.18 mm window, we need to take into account the effects of clouds and surface topography. The overlying clouds modulate the radiation mostly by scattering. The lower atmosphere of Venus absorbs the radiation emitted by the surface and also emits thermal radiation, since the opacity of the thick Venus atmosphere is not negligible at 1.18 mm wavelength [e.g., Taylor et al., 1997]. The effects of atmospheric absorption and emission increase with the column density of the atmosphere, which is largely determined by the surface
Figure 3. Theoretical and observed limb darkening for the NIMS channel 3. The radiances are plotted as a function of the cosine of the emission angle. The NIMS observation is corrected for the scattered sunlight. topography owing to hydrostatic balance. For example, the atmospheric column density is smaller in a highland area, leading to smaller atmospheric absorption and emission. In addition the intensity of thermal radiation from the surface is a function of the temperature which in turn decreases with height, so that highlands are cooler and emitting less. [20] We cannot separate the clouds’ correction from the evaluation of surface emissivity, since multiple reflection between the surface and clouds affects the intensity of radiation observed above the clouds [Hashimoto and Sugita, 2003]. Here, multiple reflection refers to phenomenon where a fraction of the upward incident radiation at the bottom of cloud layer has experienced reflections between clouds and the planetary surface more than once. This effect has a significant influence on the observed radiance under the Venus conditions, where the reflectivity of overlying clouds is high. Although multiple reflection also occurs on Earth and Mars, the degree of its effect is very small. Since the atmospheric absorption affects the multiple reflection, the correction for modulation induced by clouds is also connected to the surface topography via atmospheric opacity. [21] To estimate the cloud properties we follow the method of Carlson et al. [1993b] which used NIMS channels 5 and 7 (1.74 and 2.31 mm, respectively). At these wavelengths, visible features are almost entirely due to clouds, which modulate the emission from the lower atmosphere, with no significant surface contribution [e.g., Taylor et al., 1997]. It is not necessary to take into account the effect of multiple reflection at these channels, since photons reflected by clouds are absorbed by lower atmosphere (Figure 2). Using the NIMS channel 5 and 7, we can estimate the cloud properties without the influence from the surface emissivity. [22] Then, we calculate the upward radiance at the top of the atmosphere in channel 3 as a function of surface emissivity. The cloud properties estimated from the NIMS channels 5 and 7 and the surface topography are input to the radiative transfer model. The surface emissivity is evaluated in accordance with the computed relation between
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the surface emissivity and the radiance at the top of the atmosphere. 2.6.1. Evaluation of Cloud Properties [23] The clouds of Venus have been observed by remote sensing as well as in-situ measurement techniques [e.g., Esposito et al., 1983]. The observations indicate that the clouds of Venus are vertically stratified with three layers, which are called the upper, middle, and lower clouds, respectively [e.g., Knollenberg and Hunten, 1980]. These clouds are made up of a few different particle size components, which are often referred with mode 1, 2, 20, and 3 [e.g., Knollenberg and Hunten, 1980; Pollack et al., 1993]. The upper and middle clouds are generally uniform and featureless, while considerable spatial variations in cloud properties are observed in the lower cloud [Marov et al., 1980; Ragent and Blamont, 1980]. [24] We start with a cloud model developed by Pollack et al. [1993] and assume that the spatial variations in cloud properties are caused by variations in the number densities of mode 20 and 3 particles in the lower clouds. Carlson et al. [1993b] demonstrated that the NIMS observation is well reproduced by varying the number densities of mode 20 and 3 particles. Their results are verified by our radiative transfer model (Figure 4). Since the wavelength dependence of scattering varies with cloud particle size, channels 5 and 7 enable us to distinguish the variation in the abundance of mode 20 and 3 particles (Figure 5). 2.6.2. Surface Topography and Atmospheric Structure [25] The surface topography of Venus is measured by the Magellan spacecraft with vertical accuracy of about 50 m [Pettengill et al., 1991]. We assume that the Venusian lower atmosphere is horizontally uniform and the vertical temperature profile is given by the Venus International Reference Atmosphere (VIRA) [Seiff et al., 1985]. Below the zero altitude (6052.0 km from the center of Venus), the temperature lapse rate in the lowermost layer is extrapolated. For
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Figure 5. Venus cloud maps. (a) A map of the optical thickness of lower clouds. (b) A map of the modal radius of lower cloud particles. pressure extrapolation, the hydrostatic pressure relation is integrated using the temperature profile. We also assume that the temperature of the planetary surface equals to the temperature of atmosphere in contact with it, since the temporal variation of the temperature in the Venusian lower atmosphere is expected to be extremely small [e.g., Seiff, 1983]. 2.6.3. Emissivity Estimate and Spatial Smoothing [26] Using our radiative transfer model, we can estimate the surface emissivity based on the band-3 radiance which has been corrected for the effects of scattered day-side light and limb darkening. However, because the band-3 radiance depends only weakly on the surface emissivity due to multiple reflection between clouds and planetary surface [Hashimoto and Sugita, 2003], small errors in both observation and correction lead to a large error in surface emissivity estimate. In fact, not all the data points in the band-3 radiance have values that correspond to surface emissivity values between 0 and 1. Thus, we calculated the surface emissivity beyond the physically significant range (i.e., 0 through 1) by extrapolating the results of radiative transfer model and discarding data points whose is smaller than 1 or larger than 2. Such a criterion for data selection increases the available data and helps us to obtain a sufficient sample to infer the Venus surface composition. We reduced the random noise by averaging the data within a circle of radius 250 km. This smoothing procedure should not change general trends of the data since the spatial resolution of the Venus surface observation is limited to about 100 km owing to cloud blurring [Hashimoto and Imamura, 2001].
3. Results and Discussion Figure 4. Scatterplot of the NIMS channel 5 against the channel 7. The NIMS data are corrected for the effect of limb darkening. Six theoretical curves are also shown. For each curve, the ratio of mode 3 to mode 20 content is constant, and the total opacity of the lower cloud is varied. The lowermost curve is for a cloud that contains no mode 3 particles, while the uppermost curve is for a cloud in which the lower cloud is thoroughly composed of mode 3 particles.
3.1. Surface Emissivity [27] The estimated surface emissivity at 1.18 mm window wavelength is shown in Figure 6. Although the effect of scattered sunlight increases toward the terminator, our result does not show such a bias. Also there is no noticeable tendency varying with the emission angle, while the effect of limb darkening depends on the emission angle. These facts indicate that both the corrections for the scattered sunlight and limb darkening work fairly well.
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erties. However, the Northern Hemisphere does not exhibit such a behavior associated with the modal radius in the lower clouds. [33] The good fit of Carlson’s method (Figure 4) strongly suggests that it works well for evaluating the cloud properties. We did not identify any apparent error in the estimation of surface emissivity, though there may be a problem with the correction for the effect of clouds. It would be an important future work to develop a declouding method that uses the spectra of 1.73 and 2.3 mm window. Some instruments aboard the Venus Express spacecraft are observing the spectra of 1.73 and 2.3 mm window [e.g., Baines et al., 2006], though the Galileo NIMS did not observed them in a mapping mode [Carlson and Taylor, 1993]. Figure 6. Surface emissivity at 1.18 mm window wavelength. (a) A map of surface emissivity. (b) A map of surface emissivity that is averaged for a region of a circle of radius 250 km. [28] However, the mean surface emissivity in the Southern Hemisphere is greater than that in the Northern Hemisphere. This feature may be an artifact due to incorrect evaluation of cloud properties, since the modal radius of lower clouds also exhibit such behavior (Figure 5b). The result of declouding that is a correction for this effect is discussed further in section 3.2. [29] Although the surface emissivity map exhibits a considerable amount of noise, there is a substantial regional variation. For example, Ishtar Terra, Eistla Regio, and Alpha Regio have relatively low emissivity. In contrast, Bell Regio and a band of region from Tahmina Planitia to Fonueha Planitia have higher emissivity values. We will not further discuss a regional difference in emissivity, since surface emissivity estimation might be affected by overlying clouds. 3.2. Declouding [30] We created a declouded image corrected for cloudinduced contrast (Figure 7a). The cloud-induced contrast was corrected on the basis of the cloud properties evaluated by the method by Carlson et al. [1993b]. Although the cloud-induced contrast depends on the surface emissivity, we assumed that surface emissivity is uniform ( = 0.85) during this declouding procedure. [31] It is clearly shown that all the major features in the topography can be recognized in the declouded image, even though the declouding process is apparently a conspicuous noise source. For comparison, the thermal emission at the NIMS channel 3 wavelength was computed on the basis of the Magellan surface topography, assuming no clouds and uniform surface emissivity (Figure 7b). The correspondence between the declouded image (Figure 7a) and the synthesized image (Figure 7b) indicates that our declouding procedure effectively removes the influence of clouds. [32] The declouded image shows a north-south asymmetry that is not noticeable before the declouding. Since there is also a hemispheric asymmetry in the modal radius of lower clouds, their appearance of high surface emissivity in the southern hemisphere, especially the region along the limb, might be attributed to insufficient evaluation of cloud prop-
3.3. Temperature in the Lower Atmosphere [34] The intensity of leaking radiation at the 1.18 mm window depends not only on the surface emissivity but also on the temperature of lower atmosphere [e.g., Taylor et al., 1997]. A deviation from the assumed temperature profile will cause an error in the surface emissivity estimation. To obtain a rough estimate of the sensitivity to the temperature variation, we computed the intensities of the leaking radiation with varying the temperature of the lower atmosphere. The range of temperature deviation from the VIRA profile is roughly estimated to be about ±2 K, if the spatial variation in the declouded NIMS channel 3 image is entirely attributed to the deviation in the atmospheric temperature. [35] In other words, our analysis of the NIMS channel 3 indicates that horizontal temperature variation in the lower atmosphere is no more than ±2 K. This is consistent with results from observational and theoretical studies. While measurement accuracies are no less than ±4 K, temperature data from the Venera probes indicated that surface temperatures scatter over no more than a few K [Seiff et al., 1985]. Extrapolation of the temperature profile measured by Pioneer probes indicates surface temperatures from 731 to 735 K [Seiff et al., 1985]. Theoretical study also indicated that variations of temperature in the deep atmosphere is as
Figure 7. Thermal emission at 1.18 mm window wavelength from the surface and the lower atmosphere. (a) A declouded image that is corrected the cloud-induced contrast. (b) A synthesized image based on the Magellan topographic map.
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small as 0.1 K, because of the large heat capacity of the deep atmosphere [Stone, 1975]. [36] The NIMS channel 3 image also shows no clear evidence for a variation in temperature in the deep atmosphere with latitude or with local time; temperatures in the deep atmosphere are almost constant in spite of the variation in the insolation as a function of distance from the subsolar point. There might be a small hemispheric asymmetry in the temperature of deep atmosphere causing the north-south asymmetry in the declouded image, though that asymmetry would be attributed to the declouding process. A further study is needed to evaluate the actual horizontal variation in the temperature of deep atmosphere and its influence on the circulation of Venusian atmosphere. 3.4. H2O Abundance [37] In the estimation of surface emissivity, we assumed that H2O mixing ratio is 30 ppmv and is constant below the cloud deck. Since water vapor absorbs some radiation at 1.18 mm window, inhomogeneous distribution of water vapor causes an error in surface emissivity estimation. We computed the intensities of the leaking radiation with varying H2O mixing ratio, and found that variation of about ±10 ppm in H2O mixing ratio is required to reproduce the observed contrast in the declouded image. [38] However, it is unlikely that variation in the H2O mixing ratio is a major contributer to the spatial variation in the NIMS channel 3 image. Spatial variation of the H2O abundance in the lower atmosphere had been evaluated by Drossart et al. [1993]. They searched for its spatial variations from the analysis of the NIMS complete spectra at maximum spectral resolution using the 1.18 mm window. The result of their analysis is that the water vapor abundance shows no horizontal variation exceeding 20% over a wide latitude range on the nightside of Venus. Hence, the variation in the H2O mixing ratio is not enough to explain the spatial variation in the NIMS channel 3 image. 3.5. Emissivity and Altitude [39] Probability densities of surface emissivity (after smoothing) for three different altitude ranges are shown in Figure 8. We can see a general trend of negative correlation between emissivity and altitude; the emissivity of lowlands is generally higher than that of highlands. Although it is difficult to discuss the difference in emissivity in local scales (e.g., a few hundred km) due to noise emerging from the declouding procedure, the majority of lowlands have higher emissivity than the majority of highlands. Whereas there is an uncertainty in the results due to the not well known parameters of the atmosphere, the difference of emisivity between lowlands and highlands is more than 0.3. [40] The temperature lapse rate of the lowermost atmosphere can cause an altitude-dependent bias in the surface emissivity estimation. When a larger temperature lapse rate was used in the surface emissivity estimation, emissivities of highlands and lowlands are estimated to be higher and lower, respectively. If we use the temperature profile in which the lapse rate of the lowermost atmosphere is about 1 K/km larger than that of VIRA, the difference in the estimated surface emissivity between lowlands and highlands will disappear. However, such a temperature profile
Figure 8. Surface emissivity distribution for three different altitude ranges. Solid curve is for lowlands (z < 0 km); dashed curve is for intermediate altitudes (0 km < z < 2 km); and dotted curve is for highlands (z > 2 km). is statically unstable compared with the adiabatic lapse rate [Seiff et al., 1985]. It is unlikely that unstable stratification is observed all over the nightside of Venus, since there is no heat source which maintains the superadiabatic lapse rate. [41] It is also unlikely that horizontal variations in the H2O abundance and/or atmospheric temperature cause an altitude-dependent bias in the surface emissivity estimation. If H2O abundance and/or atmospheric temperature vary in association with surface topography, the estimated emissivity might have a bias related to the altitude. However, no observation indicates that the variation in the H2O abundance or atmospheric temperature correlates with the surface topography. Thus, it is likely that there is a difference in surface emissivity between lowlands and highlands. [42] Although we will not argue a regional difference in emissivity, a regional analysis will be useful to test the validity of emissivity estimate independent from any uncertainties in the retrieval. For example, there is no reason to expect that the volcanoes and coronae would be granitic, particularly as they are geologically young (erupted during today’s dry Venus). It would be useful to plot the emissivity of the tesserae separately from the other highlands to see if they are similar. Also, a study on the correlation of morphology and retrieved emissivity would be interesting, since morphology of many volcanic structures allows us to infer the chemical composition of the lavas [e.g., Head et al., 1992]. Using the gravity data, we can examine whether there is a systematic difference in emissivity in the same altitude region between isostatically compensated and dynamically supported highlands [e.g., Smrekar and Phillips, 1991]. [43] It is worth noting that material properties such as grain size could alter the emissivity, which may correspond to surface roughness or age (older terrains may acquire impact debris over time). It is also possible some differences might be due to weathering or that older terrains are covered by more impact debris which subdues the spectral absorptions. Chemical weathering is a function of altitude, and might be important in connection with anomalous radar
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reflectivity [e.g., Klose et al., 1992; Schaefer and Fegley, 2004]. These issues will require further consideration. 3.6. CO2 Continuum Absorption [44] The opacities of pressure-induced continuum absorption are important and poorly known factor for calculating the emission from the lower levels of the Venus atmosphere. We used the cloud model of Pollack et al. [1993] to determine the continuum absorption coefficients, but these values are somewhat different from those used in other analyses [e.g., Be´zard et al., 1990; Pollack et al., 1993]. The emissivity contrast between highlands and lowlands revealed by our analysis will decrease with decreasing the continuum opacity. Since more radiation is absorbed above lowlands than above highlands, the decrease in the continuum opacity allows lowlands to emit more radiation. Therefore, smaller continuum opacity leads to an estimation of smaller emissivity in the lowlands. There is still a large uncertainty in the parameters of the clouds, and ambiguity between average cloud opacity and continuum opacity. If the total cloud opacity is larger than that of the Pollack’s cloud model, the emissivity difference between highlands and lowlands retrieved by our analysis may decrease or even disappear. [45] To analyze the thermal emission from the Venus nightside, Meadows and Crisp [1996] used a different approach for modeling the CO2 far wings, instead of a constant binary continuum absorption coefficient. They determined the coefficients of a sub-Lorentzian profile by fitting laboratory data. Assuming constant surface emissivity, they derived that temperature lapse rate is dynamically stable (subadiabatic) as opposed to the dynamically unstable (superadiabatic) lapse rate found in this study. If temperature of the Venusian lower atmosphere is horizontally uniform, the results of Meadows and Crisp [1996] indicates a trend of emissivity increasing with altitude. The difference between their and our results indicates that modeling of the CO2 far wings is crucial to sense the surface and lower atmosphere of Venus. [46] It is quite obvious that we need detailed modeling and more accurate laboratory measurements for pressureinduced continuum absorption. Although the approach of Meadows and Crisp [1996] is no less empirical than the constant binary coefficient, it may provide insights into the processes responsible for the CO2 continuum absorption. There is also certainly a need for constraining the total cloud opacity. Either by new in situ measurements or possibly by spectroscopy with higher spectral resolution, more reliable estimates of the Venusian surface and lower atmosphere will be obtained.
4. Implication for Venus Geology and Evolution [47] Near-infrared emissivity can be used to discriminate the rock compositions [e.g., Ross et al., 1969; Baird, 1984b]. The emissivity, , is related to the reflectivity, r, by Kirchhoff’s law ( = 1 r). The reflectance spectra at near-infrared wavelength is most sensitive to the presence of iron-bearing mafic minerals [e.g., Burns, 1993]. For mafic rocks with large amounts of mafic minerals, the emissivity around 1 mm is high; while for felsic rocks with small amounts of mafic minerals, the emissivity around 1 mm is low [e.g., Baird, 1984a].
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[48] Our results indicate that there is a significant difference in emissivity between highlands and lowlands. The relative difference in surface emissivity between lowlands and highlands are consistent with lowlands materials being composed of generally mafic rocks, whereas highlands materials are composed of generally felsic rocks. We cannot yet determine which value of corresponds to which types of rocks, such as basalt, andesite, and granite. Nevertheless, the observed large disparity in between the lowland and highland is very difficult to be accounted for by the emissivity variation of basalts alone but requires presence of more felsic rocks, such as granites [e.g., Hashimoto and Sugita, 2003]. [49] This inference is consistent with several observations. The radar images of the Venusian surface by Magellan revealed that most landforms in the lowland plains suggest low-viscosity materials, which are characteristic of mafic magmas [e.g., Head et al., 1992]. The data of Venera and Vega landers are interpreted as that lowland materials are dominated by mafic rocks [e.g., Pieters et al., 1986; Kargel et al., 1993]. The inferred composition contrast between lowlands and highlands is also consistent with the principle of isostasy that is the application of Archimedes’ principle to the crust. Isostatic compensation at the surface can be achieved either by variation of crustal density or by variation of thickness of homogeneous crust. The estimated emissivity indicates that highland materials are less dense than lowland materials, since density of mafic rocks is higher than that of felsic rocks. [50] One interpretation for the felsic highland crust is that highlands are composed of granitic rocks. The presence of granitic rocks would imply that Venus had have an ocean and subduction in the past, since it has been suggested that remelting of oceanic crust combined with water along subduction would have caused the formation of granitic magmas [e.g., Campbell and Taylor, 1983]. If Venus had been an Earth-like planet where the surface was covered with ocean and subduction processes were working, granitic magmas would be generated like the present Earth. Although the present Venus is a dry-planet where granitic magmas would not be generated any more, ancient continents might have survived until today. [51] A wet origin for Venus is also suggested by numerical simulations of late-stage accretion which begins after the end of oligarchic growth of planetary embryos [e.g., Morbidelli et al., 2000; Raymond et al., 2004]. Such models indicated that Venus’s initial endowment of volatiles is likely similar to that of Earth. In fact, there is similarities between Venus and Earth in the near-surface inventories of carbon and nitrogen [e.g., Le´cuyer et al., 2000]. It indicates that Venus’s inventory of water was once similar to that of the Earth at the end of planetesimal accretion. In addition, loss of Venus’s hydrogen, that is necessary for early wet Venus to evolve into the current dry Venus, is suggested by the hydrogen isotopic ratio (D/H) measured in the atmosphere of Venus [e.g., DeBergh et al., 1991; Donahue and Hodges, 1992]. The factor of 100 enrichment in D/H compared to the Earth’s ocean is generally considered to be evidence that Venus started with at least 100 times as much hydrogen as it has now [Donahue et al., 1982]. Since early Sun’s luminosity is fainter than the present, early Venus
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would have avoided going into a state of runaway greenhouse [e.g., Kasting, 1988]. If Venus was endowed with sufficient water, early Venus would have had an ocean for more than a billion years [Hashimoto et al., 2007]. [52] It would be an important future work to look for other evidences of ancient ocean on Venus. Unfortunately, we cannot infer the ancient Venus from the observation of Venusian surface morphology, since the age of the Venus surface is no more than 1 billion years [e.g., Basilevsky et al., 1997, and references therein]. However, chemical signatures of water such as hydrous minerals may remain over a longer time period [e.g., Johnson and Fegley, 2000], even as the surface morphology has been erased. It is also reasonable to expect that ancient ocean has left its traces in the planetary interior, since the presence of water should have a great influence on the evolution of planets through controlling plate tectonics and mantle convection.
5. Summary [53] We evaluated the emissivity of the Venus surface at 1.18 mm wavelength using multispectral images obtained by the Galileo NIMS. Although the surface of Venus is shrouded in a thick atmosphere and clouds, our declouding procedure adequately remove the influence of clouds. Our analysis has revealed that the emissivity of lowlands is generally higher than that of highlands. This observation indicates that highland materials are generally composed of felsic rocks, while lowland materials are dominated by mafic rocks. Such a composition contrast between lowlands and highlands is consistent with the principle of isostasy. Since formation of granitic magmas is likely related to water, there is an implication of ancient wet Venus in the presence of granitic terrains. [54] It would be an important future work to confirm the presence of granitic rocks with the intention of revealing the history of the Venus surface environment. There are five spectral windows in the Venus atmosphere between 0.85 and 1.18 mm that are sensitive to the surface property. We will be able to evaluate the surface emissivity at these spectral windows in the same way as described in this paper. Such spectral information will help us estimate the mineralogic composition of the Venus surface. [55] It is also worth mentioning that the declouding procedure used in this study is also useful in retrieving temperature profiles in the Venus lower atmosphere. There are several spectral windows in the near-infrared wavelengths that can sense temperatures at several altitude levels below the cloud top. A multilevel global temperature fields retrieved from multiple-wavelength images will provide new insights into the general circulation of the Venus atmosphere. [56] Acknowledgments. The authors express their thanks to anonymous reviewers for their constructive comments on this manuscript. G.L.H. would like to thank Teruyuki Nakajima for detailed discussion of radiative transfer modeling, Naomoto Iwagami for helpful discussion of gas absorption, Yutaka Abe for fruitful discussion of early evolution of terrestrial planets, Yukari Tsutsumi for her help in developing computational code, and Constantine Tsang for his technical aid in preparing figures. G.L.H. and S.S. thank Grants-in-Aid for Scientific Research and The 21st Century COE Program of Origin and Evolution of Planetary Systems of Ministry of Education, Culture, Sports, Science and Technology (MEXT) for support. M.R.S. is supported by Portuguese FCT project PDCTE/FNU/49822/2003 and FCT scholarship BSAB/584/2006.
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K. H. Baines, R. W. Carlson, and L. W. Kamp, Jet Propulsion Laboratory, California Institute of Technology, M/S 183-601, 4800 Oak Grove Drive, Pasadena, CA 91109, USA. (
[email protected]; rcarlson@lively. jpl.nasa.gov;
[email protected]) M. S. Gilmore, Department of Earth and Environmental Sciences, Wesleyan University, 265 Church Street, Middletown, CT 06459, USA. (
[email protected]) G. L. Hashimoto, Laboratory for Earth and Planetary Atmospheric Science, Organization of Advanced Science and Technology, Kobe University, 1-1 Rokkodai, Nada, Kobe 657-8501, Japan. (
[email protected]) M. Roos-Serote, Lisbon Astronomical Observatory, Tapada da Ajuda, P-1349-018, Lisbon, Portugal. (
[email protected]) S. Sugita, Department of Complexity Science and Engineering, Graduate School of Frontier Science, University of Tokyo, 5-1-5 Kashiwanoha, Kashiwa, Chiba, 277-8561, Japan. (
[email protected])
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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114, E00B40, doi:10.1029/2008JE003316, 2009
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Climate evolution of Venus F. Taylor1 and D. Grinspoon2 Received 19 December 2008; revised 5 March 2009; accepted 16 March 2009; published 20 May 2009.
[1] The processes in the atmosphere, interior, surface, and near-space environment that
together maintain the climate on Venus are examined from the specific point of view of the advances that are possible with new data from Venus Express and improved evolutionary climate models. Particular difficulties, opportunities, and prospects for the next generation of missions to Venus are also discussed. Citation: Taylor, F., and D. Grinspoon (2009), Climate evolution of Venus, J. Geophys. Res., 114, E00B40, doi:10.1029/2008JE003316.
1. Introduction [2] A dramatically hot and dry climate is found on Venus at the present time, probably the result of the atmosphere following a divergent evolutionary path from a more Earthlike beginning. A key question is to what extent the two planets and their early inventories of gases and volatiles were physically alike, given their common origin in the same region of the young Solar System. Venus and Earth have nearly the same size and mass, and most current models suggest they have similar chemical compositions and interior structures. However, factors such as the small discrepancy in mean density (after allowing for compressional effects, see Ringwood and Anderson [1977]), and the absence of an internal dynamo [Luhmann and Russell, 1997], as well as discrepancies in the abundances of the noble gases [Pepin, 2006], fuel a lively debate about the extent to which the two planets can be considered to be a close match owing to essentially identical origins. The common assumption of identical origins is also clouded by the possibility of stochastic variations in late accretion history leading to unequal volatile inventories [Morbidelli et al., 2000] or volatile loss and interior processing through catastrophic early impacts [Davies, 2008; Zahnle, 2006; Alemi and Stevenson, 2006]. Even if we understood these issues, deriving the path and time scales of Venus’ divergent evolution to its present state would still present numerous challenges, not least in terms of explaining the high temperature and pressure and low water abundance at the surface. [3] Taylor et al. [2007, p. 160] posed a thought experiment: Suppose Venus and Earth had been swapped at birth – that is, at the time when they had accumulated virtually all of their present mass but before their atmospheres were fully evolved. Venus, with its slow retrograde rotation would then be at one astronomical unit from the Sun, and Earth somewhat closer. Venus would still have any bulk compositional differences it may have acquired as a result of forming at the closer position to the center of the protosolar cloud. What would this part of the Solar System look like today?
1 2
Department of Physics, Oxford University, Oxford, UK. Denver Museum of Nature and Science, Denver, Colorado, USA.
Copyright 2009 by the American Geophysical Union. 0148-0227/09/2008JE003316$09.00
In particular, would the atmospheres of Earth and Venus still be as dramatically different in composition, surface temperature and pressure? Would a repositioned Earth, its climate provoked by twice as much solar irradiance and its atmosphere exposed to a more aggressive solar wind, have developed a climate like that of present-day Venus? Would Venus, left more to its own devices at 1 AU from the Sun, be the planet with plate tectonics and oceans, and the one whose inhabitants were conducting our current explorations? To what extent might the chance formation of Earth’s singular moon have broken any potential symmetry? [4] In seeking to understand questions like this about the formation and initial states of the two planets, and how and by what stages their atmospheres progressed to the presently observed conditions, considerable use is made of the findings of each mission flown to Venus by spacecraft. Many of the measurements still needed (for instance, seismic determinations of the interior structure of Venus) will require substantial investment in new technologies, while other important investigations (such as the history and present activity of volcanism) can be addressed with available techniques. In this paper, we look at the current state of understanding of the atmosphere, interior, and nearspace environment of Venus, and the interactions between them that produce the climate at the surface and control its evolution (Figure 1). We review actual and potential progress on these topics, particularly that beginning to be made in the light of new results from the European Venus Express mission [Svedhem et al., 2009], and the important gaps that could be addressed by future planned projects. [5] The plan of the paper is as follows. Section 2 summarizes briefly what is known about the present-day climate, implicitly defined as the state of the atmosphere, and the main processes thought to be responsible for the energy balance and stability of the system. In section 3, some major topics relevant to possible long-term changes in the stability of the atmosphere are tackled, in particular the state of the interior of Venus, its thermal evolution and consequent resurfacing rates and outgassing into the atmosphere; surface-atmosphere chemical reactions and possible equilibrium states; and the loss processes at the top of the atmosphere, in particular as they concern the history of water on Venus. These are all very uncertain and hard to understand with the current very limited measurement base,
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However, the internal processes leading to critical climate elements like volcanism, global tectonics and magnetic field generation seem to be quite different, at least in outcome. New data from missions more advanced and challenging than anything attempted to date will be needed to address these; they are the subject of section 5.
2. Current Atmosphere and Climate
Figure 1. A conceptualization of the problem addressed in this paper and the formidable nature of its long-term goals. The properties and processes in and between the interior, surface, atmosphere, and near-space environment are all known to be important for determining the climate, in particular, the surface temperature. We want to understand these and their interrelationships not only for current conditions, but also throughout time from the formation of the planet to some future state when Venus may become more, or less, Earth-like. Each box in the diagram can be expanded into a range of different disciplines; for instance, the role of the atmosphere depends on its radiative transfer properties, on heterogeneous chemistry involving gas-liquid and gas-solid reactions, and on the circulation and dynamics. A further hidden dimension exists in the form of the same matrix for the Earth, and perhaps Mars also, and their relationship to Venus through common origins and processes. but incremental advances are made with each successful mission, and Venus Express was developed with the goal of making a key contribution. Section 4 reviews the incorporation of key processes into evolutionary climate models, which, in principle at least, are capable of simulating the present-day climate and showing how this may have developed over the history of the Solar System to date. Such models can also be used, within their limitations, to obtain some idea of how the climate may evolve in the future. A further dimension is added by our wish to understand, within the same physical framework, the important similarities and differences between this climate system and the corresponding one for the Earth (and also Mars and Titan, although we will not consider them in depth here). [6] The tools at our disposal for Venus climate modeling include not only the fruits of previous Venus missions and studies but also the methods that have been derived as part of the huge contemporary effort to predict the evolution of climate on the Earth. While the current emphasis in the latter is mainly on the very short term, decades to
