Temperature, Clouds, and Aerosols in the Terrestrial Bodies of the Solar System

  • F. MontmessinEmail author
  • A. Määttänen
Living reference work entry


This chapter is intended to provide a concise overview of the state of knowledge regarding the temperature, clouds, and aerosols of the terrestrial bodies of our Solar System, namely Mars, Venus, and Titan. These bodies are the planetary objects that most resemble the Earth. The atmosphere of each body is described in terms of composition and vertical structure. We distinguish and compare the extent of the various atmospheric compartments that form the atmospheric column, from the troposphere up to the thermosphere. The temperature structure is then presented, and the main causes known for explaining its variations on each body are listed. The specific roles of waves, radiation, as well as convection in shaping temperature profiles are then discussed. In a second part, the particulate components of these atmospheres, clouds and aerosols, are described in terms of their physical properties (composition, optical properties) and of their variability in both space and time. Mars , Venus, and Titan exhibit a remarkable variety of clouds and aerosols. Our knowledge about them has made considerable progress thanks to the success of space missions during the last two decades, while in parallel theoretical models have improved to the point that three-dimensional Global Climate Models now include the detailed physics of clouds and aerosols. As a result, it is now widely recognized that particulates play a key role in forcing the climate and the evolution of these bodies.


Mars Venus Tittan Atmospheres Solar System Composition Radiative transfer Troposphere Stratosphere Mesosphere Thermosphere Radiation Boundary layer Scale height Greenhouse Infrared Surface temperature Cloud layer Clouds Thermodynamic equilibrium Thermal behavior Tropopause Stratopause Mesopause Isothermal Gravity wave Radiative-convective equilibrium Radiative equilibrium Atmospheric waves Diurnal solar forcing Tides Breaking wave Composition Photodissociation Dust Dust storms Diabatic Mesoscale models Cloud Storms Mars Odyssey Mars Express Mars Reconnaissance Orbiter Venus Express Cassini 

The Thermal Structures of Mars, Venus, and Titan

When it comes to comparing their atmospheres , all the terrestrial bodies of the Solar system share common features. The atmospheric structure of these bodies is governed by the interplay of a variety of processes acting at different levels, depending on the composition of the said atmosphere and on the presence of particulate elements that interact in a specific way with their environment and in particular with radiation.

Basic Properties

The basic process that determines the vertical extent (henceforth the depth) of an atmospheric layer is related to the way solar radiation propagates through the atmosphere and deposits energy within. Whereas several terrestrial bodies are basically transparent to solar radiation (i.e., Earth and Mars), others have most of the incident solar energy trapped well above the surface (i.e., Venus and Titan). The radiative transfer in a given atmosphere is only controlled by the atmospheric gaseous and particulate composition, and this drives the difference of thermal behaviors between the bodies (Fig. 1). From the standpoint of composition, two different categories emerge from their comparison: Mars and Venus have both a dominant CO2 composition, while Earth and Titan harbor nitrogen as the main species. While CO2 is well known for its greenhouse effect, nitrogen is a radiatively inert gas.

The Various Atmospheric Compartments

Figure 1 shows the varying atmospheric structures of the terrestrial bodies of the Solar System. This comparison is only made in an average sense: most of the thermal profiles display substantial variability both in space and time, leading to large departures in shape and amplitude (especially on Mars) from the average behavior. It can be seen that in all cases, the structures share the same basic features: temperature decreases from the surface up to a certain altitude above which a large-scale warming or cooling trends start to appear (Fig. 1). The separation into various compartments (tropo-, strato-, meso-, and thermosphere) of an atmosphere is first and foremost dictated by the temperature and the main features of the thermal profile define the boundaries between the compartments (Table 1).
Fig. 1

A comparison of the mean temperature structures observed on all the terrestrial bodies of the Solar System. The limits of the various atmospheric compartments are shown here to allow comparison between bodies.

Table 1

Basic atmospheric parameters for the four terrestrial bodies of the Solar System.






Main species

CO2, N2

N2, O2

CO2, N2

N2, CH4

Surface pressure (bar)





Surface temperature (K)





Scale height (km) at near surface conditions





Tropopause temperature (K)





Tropopause pressure (bar)





Tropopause altitude (km)





(dry) adiabatic lapse rate (K/km)





The troposphere represents the part of the atmosphere that lies near the surface and where heat exchanges between the surface and the overlying gas take place. The “troposphere” term was first introduced by the French meteorologist Léon Teisserenc de Bort, who assembled two ancient Greek words tropos (tower) and sphaira (sphere) to refer to the rotary (convective or turbulent) motions of the air masses of the Earth atmosphere that exist above the surface. The Earth troposphere is the shallowest of all terrestrial bodies (∼15 km vs. >40 km elsewhere).

Above the troposphere, other compartments can be found: (i) the stratosphere where gas is vertically at rest, remaining “stratified” and exhibiting high static stability due to a positive lapse rate in its lower half; (ii) the mesosphere (literally the middle sphere) is generally recognized for hosting the coldest temperature of the whole atmospheric profile; and (iii) lastly the thermosphere which is the last part of the atmosphere that is physically connected to the rest of the column and that is usually marked by distinctly warmer temperatures compared to altitudes below. The thermosphere can be considered as the last atmospheric layer that separates the bulk gas from the interplanetary medium.

The ionized parts of atmospheres, ionospheres, will not be included in this discussion since their existence relies on widely different mechanisms and their presence has, to our current knowledge, only a minor role in cloud formation.

Mars’ Atmosphere

Atmospheric temperature is probably the most accurately and comprehensively characterized of all the Martian climate parameters. The Mariner and Viking missions produced the first temperature measurements from the orbit (Kliore et al. 1972; Lindal et al. 1979) and during the lander descent through the atmosphere (Seiff and Kirk 1977). Subsequently, temperature has been retrieved by all Mars missions, providing information on a nearly continuous basis for the last two decades. Temperature has been characterized from the ground up to the thermosphere (>100 km) thanks to a combination of nadir-looking and limb-staring observing modes relying on a variety of spectral ranges from UV to thermal infrared (Smith et al. 2001, 2003; Kleinböhl et al. 2009; Forget et al. 2009).

It is interesting to note that despite the tenuous nature of its atmosphere (among the four bodies, Mars exhibits the lowest surface pressure; 6 mbar on average, compared to 1 and 1.5 bars, respectively, for Earth and Titan, and 90 bars on Venus), Mars’ atmosphere shares characteristics with those of denser bodies. The solar radiation makes its way through the atmosphere since only a small fraction (< 10%) is retained by the suspended dust particles which absorb in the blue wavelength domain (explaining the reddish color of Mars). Since most of the radiation is able to reach the surface and warm it, heat can be sensibly transferred to the lowest layers, explaining the existence of a boundary layer.

The Mars boundary layer exhibits important similarities with the boundary layers found in terrestrial deserts. It is very active during daytime, with a peak of convective activity occurring around noon, and disappears at night, being progressively replaced by strongly stratified jet winds near the surface. The Mars boundary layer has a depth that ranges up to 10 km, implying that vigorous convective motions do occur over the same depth.

In the troposphere, the scale height H, which is the e-folding altitude decay for pressure, is defined as:

\( H=\frac{m_ag}{RT_a} \), with m a and T a the molecular mass and the temperature of the air is around 10–12 km, again very close to that of the Earth. Mars’ troposphere typically extends over a depth of ∼40 km. In this region, temperature decreases on average with a lapse rate of 4–5 K/km, to be compared with the ∼9.8 K/km on Earth and 10 K/km on Venus. In this dry configuration, the adiabatic lapse rate is equal to:
$$ {\Gamma}_d=\frac{g}{C_{pd}} $$
with g the gravity and C pd the specific heat of dry air at constant pressure. Most of the convective phenomena occurring on Mars are ruled by dry adiabatic conditions. Yet on Earth as well as in in some localized winter polar regions of Mars and also on Titan, convection switches into a “moist” regime subsequent to condensation. Moist convection has a high potential to intensify convection since it induces the release of latent heat (i.e., energy) in proportion of the condensing mass. In that case, the “moist” adiabatic lapse rate becomes (here given for that of water):
$$ {\Gamma}_w=g\frac{1+\frac{{\mathrm{L}}_v{q}_{sat}}{R_{sd}T}}{1+\frac{L_v^2{q}_{sat}}{C_{pd}{R}_{sw}{T}^2}} $$
where Γ w is the moist adiabatic lapse rate (K/km), L v is the heat of vaporization of the gas, R sd is the specific gas constant of dry air (= 287 J kg−1 K−1 on Earth), R sw is the specific gas constant of the condensing gas (e.g., for vapor, it equals 461.5 J kg−1 K−1), T is the temperature, and r is the mixing ratio of the condensing gas at saturation. Typically, in the case of water, Γ w ∼Γ d /2.

Unlike the Earth, Mars has no stratosphere, which is reminiscent of the absence on Mars of a dense ozone layer although ozone is otherwise widespread (Perrier et al. 2006; Montmessin et al. 2017). Instead, the temperature profile above the troposphere sustains a dominant cooling behavior up to the mesopause around 80 km, where temperature reaches a minimum (<100 K). Mars’ mesosphere is additionally strongly perturbed by large amplitude wave propagation (tides, gravity waves) such that, on an individual scale, a Martian temperature profile is first and foremost dominated by deep oscillations in the mesospheric portion (see, for instance, the ASI/MET temperature profile, Schofield et al. 1997).

The largest diurnal temperature variations take place in the thermosphere. On the dayside, CO2 molecules, the main gaseous component, absorb ultraviolet and near-infrared radiation and convert it into heat, leading to heating rates of >1000 K/day (Lòpez-Valverde et al. 1998). It is important to note that in this atmospheric portion, the gaseous medium progressively becomes collision-less and the radiative processes depend on a nonlocal thermodynamical equilibrium (NLTE) that adds complexity to the traditional radiative transfer description (Yelle 1991; Lòpez-Valverde et al. 1998).

Venus’ Atmosphere

Venus temperature profiles have been characterized by a variety of means: below 40 km, the atmosphere becomes so dense that traditional remote-sensing techniques become ineffective. In that part, only in situ measurements from the Venera and Pioneer probes were able to collect data (Marov et al. 1983; Seiff et al. 1980). Above 40 km, a rather complete climatology for temperature has been assembled up until 150 km and for most latitudes (see, e.g., Limaye et al. 2017). Our knowledge of the Venusian temperature profiles has been assembled into an empirical model: the VIRA (Venus International Reference Atmosphere, Seiff et al. 1985) was assembled after the Venus missions of the 1970s. These profiles described the atmospheric structure for five different latitude ranges from the equator to poles. Zasova et al. (2006) published an improved version, the VIRA-2, where supplementary data acquired after the publication of VIRA was added to the database.

Of all terrestrial bodies, Venus has the deepest troposphere since it extends up to roughly 65 km (∼up to the upper cloud layer). This fact is explainable by the very high surface pressure (>90 bars) and a pressure scale height augmented by the high temperature encountered below 50 km (at the surface, temperature reaches 700 K) that is caused by the well-known Venus’ greenhouse effect. At the high pressure of the near-surface atmosphere, CO2, the main atmospheric component, can no longer be considered as an ideal gas and actually becomes a supercritical fluid that is particularly effective at extracting heat from the surface. Unlike Mars and Earth though, the heat transfer has a more complex behavior. Most of solar radiation is deposited above 60 km (Crisp 1986) where an intense photochemical cycle is known to unfold. The major heat reservoir of the atmosphere is actually maintained by the CO2 molecules emitting infrared emission, producing in turn the high surface temperature . Within the troposphere, below the cloud layer , the temperature lapse rate is about −10 K/km. Globally, the tropospheric temperature profiles show very little variability (less than 5 K) in latitude and in local time up to 40 km, but above this level, within and above the clouds , temporal variations can be large.

Like on Mars, Venus has no stratosphere which might be explained by the lack of a discrete layer of a radiatively active species such as the one formed by O3 on Earth at 35 km. The region above the tropopause (∼65 km – 100 mbar) up to ∼120 km (1 μbar) is the mesosphere, above which the thermosphere is found. The night side upper mesosphere and thermosphere of Venus are also the source of nonlocal thermodynamic equilibrium (NLTE) emissions of CO2 and NO molecules, which are responsible for the low temperature of the nightside thermosphere (Drossart et al. 2007).

Titan’s Atmosphere

Titan’s temperature has been profiled from the orbit by Voyager and Cassini (Lindal et al. 1983; Schinder et al. 2011) and in situ by the Huygens probe (Fulchignoni et al. 2005). Titan’s thermal structure is remarkably uniform and stable compared to the other terrestrial planets. The six temperature profiles measured at various latitudes show only modest deviations (less than 1 K) between them. This uniformity observed for the atmospheric thermal behavior is also reflected in the surface temperature. The latter does not vary by more than 4 K from pole to equator. The lack of large variations is the consequence of the dense atmosphere and of the low solar forcing, resulting in a long radiative timescale.

As on all other bodies, Titan’s nitrogen atmosphere exchanges heat with the surface, creating a troposphere up to ∼45 km (0.1 bar) where temperature gradually decreases from ∼95 K at the surface on average to ∼70 K at the tropopause . Titan’s atmosphere has a substantial fraction of methane (> 1%), which, contrarily to nitrogen, is radiatively active. For this reason and because methane is also the source of a widespread organic haze, 90% of the solar radiation is trapped within the atmosphere by hydrocarbons in gaseous and particulate form (McKay et al. 1989). The result of this absorption is the presence of a stratosphere above the troposphere that extends up to > 300 km (< 0.1 mbar) (Strobel 2009) and whose lower part maintains a convective equilibrium. Above the stratopause , the mesosphere extends up until > 1000 km, although a mesopause is hard to specify on Titan. The mesosphere is almost isothermal and only perturbed by a large amplitude gravity wave (10 K).

Shaping a Temperature Profile: The Role of Radiative-Convective Equilibrium

Let us first address the concept of radiative equilibrium for an atmosphere. In such configuration, the atmosphere and the surface remain at equilibrium in the absence of other processes. The thermal profile obtained with a pure radiative equilibrium approximation pose several inconsistencies: it is too warm near the surface, too cold near the tropopause, and decays too rapidly in the troposphere. This kind of equilibrium is rather encountered in the stratospheres and the mesospheres. In fact, when solar energy is deposited at the surface, a transfer of heat from the surface to the near-surface atmospheric layers takes place. Further extraction of heat from the lowest layers to higher altitude is then needed in order to counteract the disequilibrium created by warm air accumulating near the surface. This is the concept of a radiative-(dry)convective equilibrium, where the instability created by radiation is suppressed by convective motions. This kind of equilibrium is neutrally stable as a displaced air parcel is naturally restored to its original position. This restoration mechanism is typically the one ruling the tropospheres of all terrestrial bodies. On the Earth, Titan, and potentially on Mars, latent heat release during condensation can overwhelm the dry convection mechanism, triggering moist convection phenomena that are responsible for explosive (storm-type) air parcel uplifts and cooling (e.g., convective towers).

Shaping a Temperature Profile: The Role of Waves

The presence of atmospheric waves has been reported on every terrestrial body. They are generally identified as local, large-amplitude departures from a radiative-convective equilibrated profile. Their main source generally lies in the lower layers of the atmosphere. While propagating upwards, wave amplitude increases with altitude in proportion with the related reduction of density. Wave signature can be so pronounced in certain cases that it can dominate the shape of the temperature profile. This is especially true for Mars’ mesosphere and Venus’ thermosphere where day-to-day variability of temperature variations larger than 20 K is observed.

A wave can be produced as the result of an oscillating forcing, such as the diurnal solar forcing on a rotating planet that gives rise to planetary-scale tides , which is especially reinforced on Mars due to the direct deposition of heat in the atmosphere through absorption of solar radiation by dust aerosols. Waves can also result from local imbalances (surface heating, topographic obstacles) that are relaxed by convective motions. Waves usually propagate upwards until they encounter an air mass whose flow velocity equals that of the wave phase speed or until their amplitude reaches a critical level, at which point they break. The breaking process induces a momentum transfer from the breaking wave to the local flow such that in certain circumstances, breaking waves affect the dynamics of the upper atmosphere (e.g., Earth polar mesosphere, some hints of a similar process might exist on Mars – Medvedev et al. 2011; Terada et al. 2017).

Shaping a Temperature Profile: The Upper Atmosphere

The homosphere is the part of the atmosphere where species abundance is maintained uniform in composition by mixing processes (at all scales, from local turbulent motions to large-scale advection). The homosphere comprises the layers from the surface up to the lower thermosphere (typically from 0 km to 100 km on Earth, Mars, and Venus). In this part, gases with substantial lifetime (>1 year) have their concentration remaining roughly uniform (except volatiles that can be affected by vapor pressure effects). Above the homosphere mass-dependent processes create a countereffect to mixing. Composition starts to change and each species follows a specific path as it is affected by molecular diffusion and by more efficient photodissociation processes due to a higher exposition to UV radiation. The medium becomes low enough in density that the thermalization process takes too long: radiative properties of absorbing species are no longer dictated by the kinetic temperature, and this characterizes the conditions of an NLTE regime. The radiative processes in the NLTE regime, which rely on complex theoretical considerations, have been demonstrated to apply a major imprint on the thermal behavior of an upper atmosphere such as on Mars (Lòpez-Valverde et al. 1998).

Clouds and Aerosols on Mars, Venus, and Titan

All the terrestrial bodies exhibit the presence of particulate matter in the atmosphere. Aerosols and clouds can form in various ways (see Table 2) and the differences found between the terrestrial atmospheres are also found in the cloud and aerosol characteristics (Table 2).
Table 2

Main cloud and aerosol characteristics for the four terrestrial bodies of the Solar System. Note the absence of aerosols in the case of Venus although the presence of “dry” particles in the upper haze and above has been advocated.











(H2SO4, H2O)

Various types


Mineral dust H2O, CO2


CH4, C2H6, HCN, etc.

Altitude range (km)




[45, 85]

[0, 35]

[0, 85]

[0, 70]

[0, 90]

[0, 500]

[1, 90]

Global coverage (%)











Visible optical depth (unit less)








[2, 30]

> 1

[1, 5]

Particle size (μm)





[0.1, 20]

[0.1, 1000]


[2, 30]



Particle abundance (#/cm3)











Forming mechanism



N/A photochemistry

Various sources condensation

Surface lifting


Photochemistry condensation

Some basic considerations preside the formation of clouds. Cloud formation is essentially dictated by the pressure and temperature (P, T) conditions which define the particular state (solid, ice, or gas) a species is forced to adopt. Phase diagrams provide the pressure-temperature transitions between the various phases (see Fig. 2).
Fig. 2

A schematic of a typical Phase Diagram showing the temperature-dependent evolution of the boundaries between the solid, liquid, and gas phases. Note the existence of a triple point where the three phases are all in equilibrium and can coexist together. The vapor pressure curve between gas and solid is the line that links C to A, whereas the pressure curve between gas and liquid is the one that links A to B, etc. Figure extracted from the book Principles of General Chemistry (section~11.7).

In fact, the formation of a cloud is simply the result of a change in temperature or pressure that leads a gas to cross the phase boundary and to condense into either a solid (ice crystals) or a liquid (droplets). This phase transition follows the vapor pressure curve by virtue of the Clausius-Clapeyron relation which gives the slope of the tangents to this curve, such as:

$$ \frac{dP}{dT}=\frac{PL}{T^2R}1.00em \mathrm{or},\mathrm{equivalently}0.5em \frac{dP}{P}=\frac{PL}{T^2R} dT, $$

where L is the specific latent heat of the species and R is the specific gas constant. This temperature to pressure relation for a given species does in turn impose where and when a cloud can occur. Therefore, two mechanisms explain cloud formation: (1) either temperature decreases, so the temperature-dependent vapor pressure is lowered down to below the partial pressure of the condensing species or (2) if the temperature stays constant, the condensing species concentration builds up (e.g., in case of a local source) until its partial pressure exceeds the saturation vapor pressure. While consideration of the Clausius-Clapeyron relation constitutes a prerequisite to cloud formation, it is far from representing the variety of processes at work, in particular those microphysical processes which occur at the scale of the particles and decide of its evolution. These processes are nucleation, condensation/evaporation, sedimentation, coagulation, and coalescence. Discussing the details of these processes lies beyond the scope of this chapter, yet it is important to remind that cloud formation and evolution is essentially dictated by the interplay of these processes.


Mars possesses both aerosol and condensate components in its atmosphere: mineral dust and ice clouds (Fig. 3). Both have been identified since the beginning of the space exploration endeavor, but our understanding of their nature and of the links these two components establish with the rest of the Martian atmosphere remains incomplete. Both are now studied in depth as it is recognized their interaction might hold the key of the main causes of variability of the present-day Mars’ climate.
Fig. 3

A global image of Mars showing the occurrence of the water ice clouds forming the Aphelion Cloud Belt near the topographic features of the Tharsis plateau in the near equatorial region. The elongated shape of the cloud is indicative of the prevailing orientation of the regional atmospheric flow (credit NASA/JPL-Caltech-MSSS).

Dust: A Climate Ruler

Mineral dust aerosols are ubiquitous in the atmosphere of Mars. They are torn from the surface by winds through saltation process (Greeley 2002) or are lifted by small (meter-scaled) tornadoes, the dust devils. Dust can be so abundant in the atmosphere that it may visually obstruct the entire planet during episodic events named planet-encircling dust storms (or global dust storms, see Fig. 4-left). Mineral dust shows a prominent signature in the thermal infrared around 9 μm that is diagnostic of a silicate component. The general consensus is that atmospheric dust is essentially globally uniform in composition and differs locally only by its size, which is reported to lie between a fraction of a μm to less than 2 μm (Pollack 1982; Korablev et al. 1993; Clancy et al. 2003). A ferric oxide fraction in the dust composition likely explains the ability of dust to absorb in the blue domain of the solar radiation spectrum. This absorbed energy is subsequently converted into heat for the atmosphere. This effect is the dominant source of diabatic heating and is the main ruler of the radiative-convective equilibrium that prevails on average in the troposphere of Mars. In other words, tropospheric temperature profiles are to first order controlled by the amount of dust in the atmosphere (see Fig. 4-right). Therefore, dust abundance is the prime climate parameter and its monitoring has been a quest pursued by nearly all missions since the beginning of Mars exploration, yielding a comprehensive characterization of its abundance and its variations (Smith 2004; Lemmon et al. 2005; Montmessin et al. 2017; Willame et al. 2017). The seasonal signature is prominent in the dust temporal variation. The latter obeys a cyclic behavior with a distinct trend for dust storm frequency to increase dramatically after the northern fall equinox (Ls = 180°: Ls stands for solar longitude, which runs from the northern hemisphere spring equinox, Ls = 0°, to 360° during a complete Martian year, through the northern hemisphere summer solstice, Ls = 90°, fall equinox, Ls = 180°, and winter solstice, Ls = 270°), generating a five- to tenfold enhancement of dust abundance globally.
Fig. 4

(Left) A global image of Mars made during the 2007 global dust storm (credit NASA/JPL-Caltech-MSSS). (Right) This plot, extracted from Forget et al. (2008), shows the results obtained from a theoretical model simulating a radiative-convective equilibrium in Mars’ atmosphere. It illustrates the effect dust particles have on the temperature profile. Even with a small amount of dust (cf. low dust case), dust absorption of solar radiation induces a diabatic warming of 10–20 K, which dominates the heating rate of the atmosphere. The effect of dust on surface temperature (greenhouse vs. cooling effect) is rather modest, if existing.

Understanding the dust seasonal cycle has long remained a central objective for Mars climate researchers (Newman et al. 2002; Basu et al. 2004; Kahre et al. 2005) and Mars global climate models (GCM) have been developed since their early use to simulate the fate of dust. To first order, dust is controlled by the general circulation and by sedimentation, and this mechanism paces a large fraction of the Mars’ climate seasonal variability. Yet, during the last decade, the importance of processes modulating dust vertical transport in the atmosphere and affecting the entire climate in return has been demonstrated. These processes have been identified thanks to refined description of dust-related processes with the introduction of mesoscale models (Rafkin et al. 2001) and of detailed microphysics. The local convective processes identified in the mesoscale models, in particular around the main Martian volcanoes and within the core of dust storms, are able to funnel dust in large amounts up to > 20 km of altitude where it is dispersed horizontally (Spiga et al. 2013). This mode of dust injection into the mid-troposphere disrupts the classical gradual decline of concentration, formalized by Conrath (1975) as the result of the equilibrium between mixing and sedimentation, that prevails in dust profiles on average. This subscale mode of injection creates elevated dust plumes appearing as detached layers like the ones reported by the Mars Climate Sounder on Mars Reconnaissance Orbiter (Heavens et al. 2011, 2015) and SPICAM on Mars Express (Määttänen et al. 2013). The other process affecting dust vertical distribution results from interactions with ice clouds (clouds are presented in the following). Indeed, dust particles supply a substrate for volatile species such as water or carbon dioxide to condense onto. In turn, newly formed cloud particles carry their dust seed downward as the increase of particle size subsequent to condensation implies faster sedimentation, a process known as scavenging. Scavenging is the likely cause for the water vapor observed in a state of high supersaturation (Maltagliati et al. 2011; Navarro et al. 2014). Once scavenged, dust can no longer supply a substrate to condensation, and gaseous water vapor can evolve in excess of saturation.

The Nature of Martian Cloud Formation

Clouds have been known to exist on Mars since the early days of the history of ground-based telescopic observations (Dolfus 1957; Slipher 1962). Clouds, just like dust, have been monitored by all missions since Mars Global Surveyor (MGS) that retrieved atmospheric water ice opacity with the TES instrument using the diagnostic 12 μm feature of water ice in the infrared.

On any given day on Mars, a cloud forms somewhere. It can be because humid air is advected to a colder region, or because the same air mass is forced to be uplifted by a volcano and thus to be cooled adiabatically. It can also be because some large amplitude thermal tide is propagating to an altitude where water vapor saturates. In fact, many causes govern cloud formation on Mars (and on other bodies, like Earth and Titan), and only two species are involved: most of the time water vapor condenses into ice (the pressure on Mars lies just below the triple-point of water, allowing water to exist only in vapor and ice phases) and creates micrometer-sized suspended crystals that evolve as surface fogs and higher up as stratiform clouds. At other times and/or other places, CO2, the main atmospheric constituent, can also condense, which accounts for a particularly rare phenomenon in the solar system and is likely unique among all the terrestrial bodies. CO2 cloud formation is not as well documented as water ice clouds yet it could be the dominant form, mass-wise, of atmospheric condensation.

The CO2 atmosphere follows a seasonal cycle that forces up to 25% of the atmospheric mass to condense in winter as frost on the ground of the polar regions and to later return in its gaseous form during spring and summer. Although very few details are known about CO2 when it condenses to the surface, it is predicted that it may be accompanied by low-lying optically thick clouds (Forget et al. 1995). Their formation might occasionally result in a stormy environment in the fall/winter polar regions where they form, since the latent heat released during these episodes is significant enough to cause major dynamical instabilities with an amount of released energy that is predicted to be comparable to Earth thunderstorms (Colaprete and Toon 2002). In these CO2 storms , ice crystals could reach sizes of the order of a fraction of millimeter, that is hundreds of times larger than the crystals typically observed for water ice clouds (Clancy et al. 2003; Wolff and Clancy 2003; Montmessin et al. 2006b).

Regarding the latter, a survey of their occurrence by Clancy et al. (2003) led to the distinction between two classes of ice particles, both lying in the micrometer range. The two classes refer to the two prominent water ice cloud formations observed every year on Mars: (i) largest particles (3–5 μm) are found in the Aphelion Cloud Belt (ACB) which stretches throughout the intertropical zone during the first half of the Martian year, peaking in intensity at aphelion season (Ls = 70°), and the ACB tends to be optically thicker around the most protruding topographical features of the equatorial region such as the Tharsis volcanoes or Valles Marineris for instance; (ii) the Polar Hoods (PH) consist of low-lying clouds with crystal radius of <2 μm that extend vertically from the surface up to 15–20 km (Montmessin et al. 2006a; Benson et al. 2010, 2011), and these clouds have been poorly characterized as they are mainly found at the seasons and places corresponding to the polar nights of both hemispheres which challenge their remote sensing. Nevertheless, PH clouds are predicted to account for the largest fraction of the water ice mass suspended in the atmosphere (Montmessin et al. 2004; Madeleine et al. 2012; Navarro et al. 2014).

Mesospheric Clouds

The mesospheric clouds on Mars, and the identification of their CO2 ice composition in certain observations (Montmessin et al. 2007), is a recent endeavor initiated by the Mars Express mission. The Martian atmosphere is extremely thin at mesospheric altitudes (pressure of μbar). Such cloud formation process lies at the frontier of space and thus at the limit of our canonical understanding of cloud-formation processes that prevail on Earth (note that mesospheric clouds also occur on Earth but at much higher pressure than on Mars).

The first observations of high-altitude mesospheric cloud layers on Mars came from limb images and radiance profiles with no information on the composition (Clancy and Sandor 1998; Clancy et al. 2007). Similar layers were observed in occultation (Montmessin et al. 2006b) and these were assigned to CO2 ice clouds subsequent to the nearby occurrence of supersaturated atmospheric pockets found in joint temperature retrievals. The discovery of the CO2 ice signature at 4.24 μm in nadir observations (Montmessin et al. 2007) initiated a hunt for mesospheric clouds (Määttänen et al. 2010; Scholten et al. 2010; McConnochie et al. 2010; Vincendon et al. 2011; Sefton-Nash et al. 2013). Still, mesospheric clouds might, at least to some extent, be the result of H2O ice crystal formation, as shown by Vincendon et al. (2011). These clouds form mainly in the tropics during northern hemisphere spring/summer and their altitudes vary between 45 and 85 km (daytime observations) and as high as 90–100 km (nighttime observations). Some observations have allowed the extraction of the CO2 ice cloud properties. The ice crystal effective radius exhibits some diurnal dependence: from 0.5 μm to 3 μm during daytime and from 80 to 120 nm during the night. While this may reveal a real diurnal evolution, it may also be the result of observing methods. The opacity of the daytime clouds can exceed 0.5 (at 1 μm).

It appears that the clouds are more prone to form in the coldest part of the mesosphere where a combination of propagating thermal tides (Gonzalez-Galindo et al. 2011) and gravity waves is able to generate temperature perturbations conducive to the CO2 ice crystal formation (Spiga et al. 2012).


The cloud system of Venus still holds on to its mysteries. This section reviews the current state of understanding of the clouds of Venus.

Cloud Zoology

Venus’ clouds , which enshroud the whole planet (Fig. 5-left), are organized in three layers, found between about 48 km and 68 km, surrounded by hazes above and below. The total optical depth of the clouds exceeds 30. The visible albedo of the clouds is very high (over 0.8), covering the planet and its surface with a veil that makes it look like a rather featureless billiard ball for the naked eye. The clouds are formed mainly of sulfuric acid solution droplets (75–95% in weight; Hansen and Arking 1971; Pollack et al. 1974; Barstow et al. 2012), but suggestions of crystalline particles and other substances in trace amounts have been derived in certain studies (Krasnopolsky 1989, 2017).
Fig. 5

(Left) A global image of Venus made by the Venus Monitoring Camera (VMC) on-board the Venus Express orbiter. This image, collected with a UV filter, shows a variety of structures (bright collar near the pole, mottled clouds near the equator, and dark stripes at mid-latitudes) that cannot be seen in the visible (credit: ESA). This appearance implies the presence of a UV absorber that creates contrasts in the upper cloud layer and whose composition and origin is still largely debated. (Right) From Knollenberg and Hunten (1980), this vertical profile of the cloud structure on Venus was collected by the Pioneer Venus in situ probe particle size spectrometer. Three distinct cloud layers were identified, all of them showing different particle size distributions and hence different processes at work in forming and maintaining these layers.

With a closer look, and particularly in the ultraviolet wavelengths (Titov et al. 2008), one can see that the clouds are not completely smooth, but show clear latitudinal structures in UV contrast that can also change rapidly. The constituent that absorbs at UV wavelengths has not been identified, but it seems to be correlated with circulation and clouds (Titov et al. 2008). The tropical region displays convective structures (Moissl et al. 2009), revealed by mottled, puffy clouds, whereas the mid-latitudes have a much more laminar, smoother structure (Markiewicz et al. 2007). The cloud top altitudes (optical depth of unity at wavelength of 1.6 μm, Ignatiev et al. 2009) in the tropical regions are high and remarkably constant (74 +/− 1 km), whereas at the poles they are lower (63–69 km) and exhibit temporal variations. Within the polar vortices, the cloud top altimetry correlates with thermal emission, but not with UV signatures, whereas UV contrasts seem often correlated with the cloud altimetry in the mid-latitudes (Ignatiev et al. 2009).

The possibility of water ice crystals at high altitudes has been evoked (O’Leary 1970), and such thin mesospheric layers may have been observed by SPICAV (Montmessin et al. 2010), but their composition remains to be confirmed.

Cloud Microphysical Properties and Processes

The particle size distributions, first measured in situ by the Pioneer Venus probe (Knollenberg and Hunten 1980, see Fig. 5-right), exhibit two modes composed of sulfuric acid droplets (mode 1: radius ∼0.2 μm; mode 2: r∼1 μm) from the upper haze (Wilquet et al. 2009; de Kok et al. 2011) to the lower cloud layer (Knollenberg and Hunten 1980). In addition, a third mode (mode 3: r∼3–4 μm), which is most probably composed of at least 50% by weight of H2SO4 (Grinspoon et al. 1993), was possibly observed in the lower cloud by the Pioneer Venus descent probe (Knollenberg and Hunten 1980). They suggested it to be crystalline, although this hypothesis was contested and explained as a misinterpretation caused by the instrument calibration (Toon et al. 1984). Modes 0 (r∼0.0065 μm) and 2′ (r∼1.4 μm) have been suggested (Tomasko et al. 1980; Pollack et al. 1980), but Grinspoon et al. (1993) showed that the mode 0 above the upper haze was not necessary. The mode 2′ can be simply interpreted as a slightly larger mode 2 found at lower altitudes. Observations of polarization of light have revealed that upper haze and cloud particles should be spherical (Hansen and Hovenier 1974; Rossi et al. 2015). This does not exclude the possibility of a small fraction of crystals among the droplets, or near-spherical glass-like amorphous ice particles (McGouldrick et al. 2012). The lower haze is composed only of very small (mode 1) particles (Knollenberg and Hunten 1980; Satoh et al. 2009) that are unlikely to be composed of sulfuric acid, which should evaporate at the high temperatures prevailing at these altitudes. The polar regions are dominated by polar vortices that seem to host particles that are distinctly different from the rest of the planet, either larger in size or differing in composition in the lower cloud (Wilson et al. 2008; Haus et al. 2013, 2014; McGouldrick and Tsang 2017). In addition, Luginin et al. (2016) showed that, when assuming a unimodal distribution in their analysis, the upper haze (70–90 km) particles were smaller in the polar regions than elsewhere on the planet.

The dominating microphysical processes at play in the Venus clouds depend on the cloud layer and are strongly coupled with the atmospheric chemistry providing the condensing vapors. The upper cloud is formed right around the photochemical production zone of sulfuric acid and is most probably the result of rapid nucleation and growth thanks to the vapor supply and efficient coagulation due to high particle number densities. The nucleation pathways are discussed in the end of the section. The process timescale analysis of James et al. (1997) showed that this upper photochemical cloud and the condensational cloud below can be treated separately because of their differing growth and vertical transport lifetimes. The lower and middle cloud layers are convective, as shown by the measured static stability profile (Seiff et al. 1980) and explained by top-down convection (McGouldrick et al. 2012) fueled by LW heating at the bottom and cooling at the top. These cloud layers are defined by a cycle of droplet growth, sedimentation, evaporation in the subsaturated zone below the clouds, and recycling of the vapors into the lower cloud by convection.

Particular Features of Venus Clouds

The upper cloud of Venus is a fairly unique phenomenon, since it is photochemically produced. The photochemical sulfuric acid vapor production at an altitude of about 65 km leads to the formation of this cloud layer. Photochemical haze formation is observed on other Solar System bodies (for example, Titan, Pluto), but on Venus photochemistry produces a condensable vapor that can directly nucleate and induce cloud droplet growth in large quantities.

The strong photochemical vapor source also does not necessarily guarantee droplet formation via homogeneous nucleation. The exact formation pathway of the cloud droplets is still not known. It has been supposed that the mode 1 particles act as condensation nuclei for the larger modes, but even though several candidate compositions have been proposed (Sx, meteor dust; see, for example: Young 1983; James et al. 1997; Gao et al. 2014), the nature of the condensation nuclei remains to be elucidated. Modeling studies on the ionization of the Venus atmosphere show that ion concentrations might be significant at the cloud formation altitudes (Michael et al. 2009; Plainaki et al. 2016), so ion-induced particle formation might be plausible. Neutral homogeneous and ion-induced particle formation models (Duplissy et al. 2016; Merikanto et al. 2016; Määttänen et al. 2017) could also be tested for Venus’ clouds.

The possible condensation nuclei, or another mysterious species related to cloud formation, might be at the origin of the contrasts observed at UV wavelengths. The absorptions observed in measured UV spectra have not been explained by any known spectral features of substances existing in the Venus’ atmosphere. Sulfur species have been proposed as an obvious explanation because of their prevalence, but no perfect match has been acquired.


Titan is enshrouded with an optically thick organic haze and is also marked by the recurring apparitions of condensate clouds (Fig. 6-left). The particulate matter suspended in the Titan’s atmosphere can be divided into two main categories: (i) organic aerosols formed out of photochemical processes taking place in the upper atmosphere, and (ii) condensate particles formed out of hydrocarbon gases (part of them being also produced in the upper atmosphere) condensing into a liquid or an icy phase. Titan features a complex mechanism controlling the fate of these suspended particles (Hörst 2017).
Fig. 6

(Left) A global composite image of Titan obtained by the Cassini mission and created out of a set of images collected at several infrared wavelengths allowing one to distinguish surface and atmospheric features though the haze (credit NASA Jet Propulsion Laboratory). (Right) Sketch illustrating the cycle of formation of organic haze particles, the tholins, in the Titan’s atmosphere (credit NASA). The main species involved are methane and nitrogen whose photochemical decomposition in the upper atmosphere (>900 km) initiates a cycle of complex hydrocarbon formation, leading to the fractal aggregate aerosols that populate the lower atmosphere.

Organic Haze

The organic haze is a ubiquitous component of Titan’s atmosphere. It opposes an optically thick layer to sunlight and thus limits the penetration of radiation, absorbing 40% of the incident solar flux (McKay et al. 1991). Haze particles are the major absorber shortward of 5 μm but are essentially transparent in the near-IR. The role of the organic haze in the energetic balance and in the dynamics of the atmosphere is now well established: it is known to deposit energy in the stratosphere while cooling the surface. The UVIS instrument onboard Cassini has revealed the presence of aerosols up to 1000 km (Liang et al. 2007). In fact, a large fraction of the haze is likely produced in the upper layers of the atmosphere (Fig. 6-right) (Tomasko 1980). The discovery of heavy ions in the thermosphere thanks to Waite et al. (2007) has led to the conclusion that an active ionospheric chemistry takes place and that haze particles are predominantly the result of aggregation of heavy ions followed by coagulation of neutral aggregates at lower altitudes. The end result is that haze particles, the tholins, progressively turn into aerosols exhibiting a fractal structure during their growth as they fall deeper into the atmosphere below 300 km. The fractal characteristic was hypothesized for the first time by West and Smith (1991) and was later confirmed by observations (Rannou et al. 1995, 2003).

Organic haze is present everywhere on Titan, and its vertical structure possesses intriguing features. In particular, the presence of a detached layer stretching nearly pole-to-pole above 200 km and exhibiting a distinct seasonal behavior has long remained an enigma. Rannou et al. (2002) proposed that this detached layer is paced by seasonal changes in the main meridional circulation, reversing its north-south orientation as the meridional wind changes direction between solstices.

Hydrocarbon Clouds

Considering the observed abundance of methane in Titan’s atmosphere and the saturation conditions that prevail for this species, it is highly probable that methane can potentially condense and form clouds. At the tropopause, the volume mixing ratio of CH4 is found to remain below its saturation value (∼2 × 10−2). Methane photolysis in the upper atmosphere fuels an active photochemistry that leads to the formation of a variety of secondary gases that consist mainly of higher order alkanes (e.g., C2H6, C3H8) and nitriles (e.g., HCN). Once formed, these species are mixed throughout the atmosphere and are able to reach altitudes where they can condense.

The same way dust plays the role of condensation nuclei for clouds on Mars, the organic aerosols of Titan’s haze supply substrates that permit condensation to occur. Although no direct observation has confirmed this hypothesis (this is true for Mars as well), it has been demonstrated experimentally that CH4 readily nucleates on tholin analogs (Curtis et al. 2008).

The discovery of clouds on Titan occurred later after the Voyager 1 survey. Tsai et al. (2012) give a summary of the main types of clouds that have been found on Titan: convective methane clouds (first reported by Griffith et al. (1998, 2000) and reported numerously since then), ethane clouds (Griffith et al. 2006; Brown et al. 2010), and high altitude nitrile cirrus clouds (Samuelson et al. 2007; de Kok et al. 2014).

As on any other terrestrial planet, the basics of cloud formation apply also on Titan: either saturation conditions are met due to dynamical or radiative cooling of an air parcel or through humidification, that is, the local build-up of a species concentration until it finally saturates and condenses. In the tropics, methane in the atmosphere is around saturation at 9–16 km and at 23–35 km and only small temperature perturbations (convection or waves) are required there to initiate cloud formation. In fact, Titan’s temperature profile reveals that air parcels below 15 km altitude are buoyant (Griffith et al. 2014) and thus prone to adiabatic cooling.

As explained by Griffith et al. (2005), the characteristics of the clouds, their height, and their evolution (to which we could add their vertical structure) have not been documented yet models have tentatively filled the knowledge gap left by observations. Indeed, the understanding of cloud formation processes has made considerable progress thanks to the use of GCMs and other three-dimensional dynamical models (Rannou et al. 2006; Hueso and Sanchez-Lavega 2006; Barth and Rafkin 2007). These models have the capability to reproduce the complex interplay between atmospheric dynamics, radiation, and cloud microphysics, which are needed to produce realistic simulations of cloud formation and evolution. With these theoretical tools, the cloud formation processes prevailing on Titan have been identified: updraft and adiabatic cooling within the Hadley circulation, radiative forcing at the summer poles, wave forcing, and volatile downward flow from the stratosphere.

Remarkably, Titan has a methane cycle resembling a lot to the Earth’s water cycle, involving elements such as liquid bodies, rainfall, storms, etc. No such resemblance is found elsewhere in the solar system, making Titan unique in that respect when compared to Venus or Mars. Although some storm activity is suspected to occur in the deep Martian polar nights (Colaprete and Toon 2002), Titan’s storms bear more similarities with terrestrial thunderclouds. These storms appear to be controlled first and foremost by solar insolation which dictates where regions of upwelling can occur, and therefore a seasonal excursion of the storm concentration is observed between the two polar regions through the equator.

The downward flow of volatiles produced in the stratosphere is the second main process for cloud formation on Titan. As they are produced higher up, these species diffuse downward and eventually encounter cold enough conditions to condense into the solid phase. These clouds have been both observed (Coustenis et al. 1999; Anderson et al. 2010; etc.) and simulated (Rannou et al. 2006; Lavvas et al. 2011). They form above 30 km up to 90 km, and exhibit spectral evidence of ice particles of a few microns in size. A variety of species has been detected involving hydrocarbons (C2H6) and nitriles (HCN, HC3N, etc.) essentially (a list can be found in Hörst 2017).


The Role of Clouds in the Climates of Terrestrial Bodies

On Mars, water vapor has been observed to subsist in a highly supersaturated state above 30 km of altitude during certain seasons (Maltagliati et al. 2011). This unexpected result was interpreted in terms of requirement for water to have dust particles around to condense onto. When dust has been progressively scavenged out by cloud sedimentation, no more dust “seeds” are available for condensation process to take place since water condensation onto itself (so-called homogeneous nucleation to be compared with the heterogeneous nucleation when dust is involved) requires supersaturated states much higher than ones observed (Michelangeli et al. 1993; Määttänen et al. 2005). This allows water vapor to propagate higher in the atmosphere, a result confirmed and shown to be even reinforced during southern summer (see Maltagliati et al. 2013), also manifested as a seasonal intrusion of unexpected amounts of hydrogen atoms above the exobase (Chaffin et al. 2014; water vapor is the main carrier of hydrogen). If clouds were to form every time a supersaturated state appears, it would preclude this upward propagation of water up to altitudes where an access to space is granted. Therefore, clouds play a definite role in the way water is kept from escaping to space on Mars.

Another climatic aspect of clouds that has been recently highlighted concerns their radiative effects. Even if not as profound and as ubiquitous as that of dust, the cloud radiative effect is revealed by models that account for the interactions between cloud particles and radiation. These models are able to reproduce faithfully the thermal behavior of the intertropical atmosphere around northern spring/summer as the result of the ACB imprinting its presence through near-IR absorption and warming between 20 km and 40 km, a result unachieved by previous model generations.

On Venus, because of the high visible albedo and high total optical depth, the clouds block sunlight very efficiently. The clouds absorb both in the visible and in the infrared, and about half of the solar flux is absorbed within the clouds. The surface of Venus receives only 2.5% of the total insolation received by the planet at the top of the atmosphere (Tomasko et al. 1980). In the infrared, emission by clouds induces a heating of the surface of the order of 100 K (Titov et al. 2006). Overall, the cooling effect of the clouds dominates: without the clouds, Venus would be even hotter.

Chemistry and clouds are also intimately linked on Venus. The sulfur cycle leads to the formation of H2SO4 that subsequently condenses into cloud droplets with H2O and is thus removed from the vapor phase, before re-evaporating when the droplets fall into the subsaturated zone below the clouds. The clouds thus act as a “lid” of H2SO4, leading to very small amounts of sulfuric acid above the clouds (as observed).

On Titan, the main cloud effect is probably related to the tropospheric storms that impose a “wet” adiabatic thermal behavior in the zone of their formation. The global cloud impact (GCI) on Titan stills remains to be evaluated but is likely to be secondary compared to the thick and ubiquitous organic haze. As such, GCI might be found to be very similar between Mars and Titan, at the opposite of Venus and with Earth lying in between.

Reappraising Terrestrial Clouds and Aerosols in the Solar System

Since the beginning of the 2000s, the exploration of terrestrial bodies has entered a golden age. By 2010, all terrestrial planets (Mars, Venus) and moon (Titan) had an orbiter flying around them: Mars Odyssey , Mars Express, and Mars Reconnaissance Orbiter around Mars; Venus Express around Venus; and Cassini around Saturn and Titan. Thanks to these probes and to their scientific payload, our view of these bodies, and in particular of their atmospheric composition and activity, has changed considerably to reach an unprecedented level of detail. It is interesting to note that clouds and aerosols are observed on all of these bodies, and that each of them exhibits a distinct behavior, implying widely different mechanisms at work.

Among the major achievements of these missions regarding clouds and aerosols is that it has been possible to establish a detailed list of condensing species in the atmosphere of Titan, while acquiring information on the nature, timing, and on the spatial variability of cloud formation. Likewise, our understanding of the Venusian sulfuric acid clouds has made considerable progress; their appearance and some of their properties have been characterized with greater precision. We can now distinguish various forms of cloud formation, like the equatorial mottled clouds that reveal the presence of convective cells underneath and that contrast with the brighter laminar appearance of the mid-latitude clouds stretching around the planet with the main zonal circulation. Yet, some major unknowns remain, in particular the nature of the UV absorber as well as details on the properties of the lowest cloud layer.

No other planet, except the Earth, has been so intensively monitored as Mars. Its main atmospheric parameters have been continuously surveyed since nearly two decades and will continue for the coming decade with the ExoMars Trace Gas Orbiter . The characterization of dust and water ice clouds has benefited immensely from this multi-annual survey: a comprehensive climatology of both clouds and dust is now available, supplying strong constraints to climate modelers.

This sets the basis for ongoing and future comparative planetology research as we dispose now of the essential information to put all the bodies in a frame of comparison regarding their clouds and aerosols. This work is essential for supporting theoretical studies and the analysis of observations of extrasolar worlds, since the cloud and aerosol aspect appears as a key element in both the data interpretation (the presence of clouds can potentially be a showstopper as it may obstruct the view of the surface) and also in driving several climatic phenomena (radiative feedbacks, etc.) for these bodies.


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Copyright information

© Springer International Publishing AG 2018

Authors and Affiliations

  1. 1.LATMOS/IPSLUVSQ Université Paris-SaclayCNRS, GuyancourtFrance

Section editors and affiliations

  • Agustín Sanchez Lavega
    • 1
  1. 1.Universidad del País VascoBilbaoSpain

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