# Vertical Structure of the Martian Atmosphere: The View from Mars Express

Marco Giuranna, S. Tellmann, Franck Montmessin, D. Belyaev, L. Trompet, M. Pätzold, P. Wolkenberg, M. A. Lopez-Valverde, O. Korablev, F. Forget

► To cite this version:

Marco Giuranna, S. Tellmann, Franck Montmessin, D. Belyaev, L. Trompet, et al. Vertical Structure of the Martian Atmosphere: The View from Mars Express. Space Science Reviews, 2025, 221 (3), pp.36. 10.1007/s11214-025-01155-9. hal-05029218

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# Vertical Structure of the Martian Atmosphere: The View from Mars Express

M. Giuranna<fnref n="1" /> Ⓒ · S. Tellmann<fnref n="2" /> · F. Montmessin<fnref n="3" /> · D. Belyaev<fnref n="4" /> · L. Trompet<fnref n="5" /> ·

M. Pätzold<fnref n="2" /> · P. Wolkenberg<fnref n="1" /> · M.A. Lopez-Valverde<fnref n="6" /> · O. Korablev<fnref n="4" /> · F. Forget<fnref n="7" />

Received: 25 October 2023 / Accepted: 26 February 2025
© The Author(s) 2025

## Abstract

Launched in 2003, the European Space Agency's Mars Express (MEX) has been orbiting Mars for 20 years and its instruments have performed continuous monitoring of the conditions in the Martian atmosphere, providing one of the most complete datasets of atmospheric parameters ever collected for Mars. This article provides an overview of the observations of the vertical structure of the Martian atmosphere performed by MEX, which led to the identification of peculiar phenomena that affect the atmospheric circulation and dynamics on different scales, from local to regional and global.

## 1 Introduction

The vertical structure of the Martian atmosphere can be ideally divided into three main layers: the lower atmosphere, the middle atmosphere, and the upper atmosphere. The lower atmosphere is the closest atmospheric layer to the planet's surface and extends to about 50 km in altitude. In this layer, temperature decreases with height, as it is driven by convective and radiative heat transfer. The temperature and thickness also vary greatly throughout the diurnal cycle. The middle atmosphere goes from 50 km to 100 km, and is roughly isothermal. The lower and middle atmosphere are also referred to as the neutral atmosphere. Mars' atmosphere above 100 km is designated as the upper atmosphere, where molecular diffusion dominates and temperature generally increases with altitude due to heating from the absorption of solar radiation in the far and extreme ultraviolet parts of the spectrum. The

<figure><img src="image_2.png" /></figure>

<footnote>M. Giuranna
marco.giuranna@iaps.inaf.it</footnote>

<footnote n="1">¹ Istituto Nazionale di Astrofisica, Istituto di Astrofisica e Planetologia Spaziali, Rome, Italy</footnote>

<footnote n="2">² Rheinisches Institut für Umweltforschung at the University of Cologne, Cologne, Germany</footnote>

<footnote n="3">³ Laboratoire Atmosphères, Observations Spatiales, CNRS, UPS, SU, Guyancourt, France</footnote>

<footnote n="4">⁴ Space Research Institute (IKI) RAS, Moscow, Russia</footnote>

<footnote n="5">⁵ Royal Belgian Institute for Space Aeronomy (BIRA-IASB), Brussels, Belgium</footnote>

<footnote n="6">⁶ Instituto de Astrofísica de Andalucía (IAA/CSIC), Granada, Spain</footnote>

<footnote n="7">⁷ Laboratoire de Météorologie Dynamique/IPSL, Sorbonne Université, ENS, PSL Research University, Ecole Polytechnique, CNRS, Paris, France</footnote>
state of a planetary atmosphere can be described by the interdependence of the temperature, pressure, and density as a function of the altitude. The knowledge of the vertical temperature profiles is essential for an understanding of the dynamical, thermodynamical, and chemical processes occurring in the atmosphere.

Knowledge of the structure of the neutral atmosphere of Mars advanced rapidly in the age of space astronomy in the late 60s and early 70s. The exploration of tropospheric temperatures started with radio occultation measurements from the flyby of the Mariner 4 spacecraft in 1964 (Kliore et al. 1965), followed by two other flybys of the Mariner 6 and 7 in 1969 (Hogan et al. 1972). Then, the successful insertion of Mariner 9 spacecraft in orbit in 1971 allowed a first “long-term” monitoring (about half a Martian year) of the atmospheric temperatures below 60 km using both radio science and thermal infrared observations (Kliore et al. 1972; Hanel et al. 1972a,b). The remote sensing experiments from the flybys and the orbiting spacecraft provided important information about the Martian atmosphere. The mean surface pressure was found to be near 5 mbar, much lower than was previously accepted and, in conjunction with ground-based spectroscopy, it was confirmed that the atmosphere is predominantly CO₂. They also indicated that the temperature structure up to 45 km altitude is highly variable, dependent on latitude, season, and the dust content of the atmosphere (Kliore et al. 1972; Hanel et al. 1972b). The major effects of the suspended dust on the temperature profiles was first displayed during the early stages of the Mariner 9 mission. The observations started during the Southern summer, also known as the “dusty season” of Mars. The vertical temperature profiles obtained from thermal infrared observations and radio occultation experiments indicated a significant atmospheric warming due to the absorption of the solar radiation by dust (Martin and Kieffer 1979), as well as a drop of atmospheric temperatures by 30 K and 40 K at 2 and at 0.3 mb, respectively, during the dissipation of a dust storm (Conrath 1975).

The temperature data collected by the Mariner experiments were limited to the lowest 35 km (occultation) and 45 km (IR sounding). The entry, descent, and landing (EDL) of the two Viking landers provided the opportunity to make in-situ measurements of Martian atmospheric structure and returned a detailed snapshot of the density, pressure, and temperature of the Martian atmosphere over a wide range of altitudes extending from the thermosphere to the lower atmosphere. In particular, the Atmosphere Structure Experiment (ASE) on the two Viking landers provided high vertical resolution profiles of temperature, pressure, and density from 120 km down to the surface at two locations of Mars (Seiff and Kirk 1977). Measurements from the Viking upper atmosphere mass spectrometer instrument (Nier and McElroy 1977) allowed profiles to be extended up to an altitude of 200 km in the case of Viking 1 and up to 175 km in the case of Viking 2 (Seiff and Kirk 1977). 21 years later, the EDL of the Mars Pathfinder spacecraft provided the first opportunity to study Mars' nighttime atmospheric structure with similar accuracy, vertical resolution, and altitude range (Magalhaes et al. 1999). These in-situ measurements (Fig. 1) complement atmospheric temperatures derived from infrared spectrometers on orbiters, which provide global coverage of the lower atmosphere over an extended period of time at lower spatial and vertical resolution, and are further complemented by meteorological measurements from landers, which provide an extended time series of high-accuracy measurements at one location at the surface. Data collected by the two Viking landers and by the Mariner 9 orbiter pointed out weak inter-annual and seasonal variability of atmospheric temperatures at 0.5 mb during low dust conditions, while day-to-day variations of atmospheric temperature were observed in high northern latitudes, likely related to a passing weather system (Santee and Crisp 1993). Night-time temperature drops by around 80 K were observed at low altitudes during the southern summer, especially at subsolar latitudes (Santee and Crisp 1993) whereas large diurnal variations of 15-30 K were consistently observed below 30 km over most of the planet (Hanel
<figure><img src="image_3.png" /><figcaption>**Fig. 1** Vertical structure of the Martian atmosphere. Temperature profiles inferred from deceleration measurements aboard the Viking 1 (blue), Viking 2 (green) and Pathfinder (red) landers. See text for more details. Modified from Haberle (2015)</figcaption></figure>

et al. 1972b). Near-surface atmospheric temperature inversions were also observed at night, as a consequence of the fast decrease in surface temperatures. The average lapse rate of the Martian troposphere was found to be ~2.5 K km⁻¹, which is much lower than the dry adiabatic lapse rate (~4.5 K km⁻¹ in the case of Mars). Consequently, Mars has quite a stable troposphere. The additional heating comes from the absorption of solar radiation by suspended dust particles, and the temperatures are further stabilized by vertical heat fluxes associated with large-scale circulation systems (Haberle et al. 2017, and references therein). Daytime boundary layer convection can extend to relatively high altitudes (~15 km) on Mars. In such regions, the lapse rates are close to the adiabatic value. Evidence for daytime convection on Mars can be seen in the Viking lander entry profiles (Fig. 1), which show a near adiabatic lapse rate between the surface and 6 km. Above 15 km, temperatures continue to decrease with height, but are controlled almost entirely by radiation rather than convection. A temperature inversion around 15 km was observed by the Pathfinder entry profile (Fig. 1). In the Martian mesosphere, temperatures become nearly constant, with superimposed oscillations due to the adiabatic heating and cooling associated with vertically propagating planetary waves which, in turn, are associated with a global system of thermal tides, planetary waves, and gravity waves. In the upper mesosphere, around 90-120 km, additional diurnal oscillations up to about 30 K are expected due to non-LTE solar heating in the near-infrared by a large set of CO₂ ro-vibrational bands (López-Valverde et al. 2000). Finally, in the thermosphere, temperatures increase because of heating due to the absorption of solar radiation in the far and extreme ultraviolet.

Nowadays, several datasets of atmospheric parameters, collected in a variety of different ways, are available for Mars. Data returned by orbiting spacecraft have been collected systematically on a global scale and provided the most complete datasets. Atmospheric temperatures and aerosol opacities have been consistently retrieved for several Martian years by remote-sensing infrared observations on various spacecraft. Mars Global Surveyor (MGS) has collected thermal infrared and radio occultation measurements from Ls~100° of MY 24 to Ls~80° of MY 27 (Hinson et al. 1999; Tyler et al. 2001; Conrath et al. 2000; Smith et al. 2001; Smith 2004). The Thermal EMission Imaging System (THEMIS) on Mars Odyssey (MO) (Smith et al. 2003) has been collecting data since Ls~330° of MY 25 (Smith 2009) while the Mars Climate Sounder (MCS) on Mars Reconnaissance Orbiter (MRO) (Kleinböhl et al. 2009) has been collecting data since Ls=111.3° of MY 28 (McCleese et al. 2010). Atmospheric temperatures retrieved from MGS, MRO, and MO have allowed for detailed analyses of the atmospheric structure, thereby constraining the global circulation of the Martian atmosphere across different seasons and multiple Martian years (Smith et al. 2001; McCleese et al. 2010; Heavens et al. 2011). These three missions have also investigated global dust events that occurred on Mars during MYs 25, 28, and 34, confirming
the significant impact of large dust quantities on atmospheric temperatures and circulation (Smith 2004, 2009; Kass et al. 2019). The Emirates Mars Mission (EMM) (Amiri et al. 2022) has recently completed its primary mission, having spent more than one Mars year observing the Martian atmosphere. The Emirates Mars Infrared Spectrometer (EMIRS) (Edwards et al. 2021), along with EMM's other instruments, has collected a vast collection of data providing both seasonal and diurnal time coverage of the entire planet (Smith et al. 2023). These datasets, combined with ever more accurate global circulation models, have allowed us to improve our understanding of the main physical processes that control the current Martian climate, circulation, and dynamics.

The Mars Express (MEx) ESA mission (Wilson and Chicarro 2004) was launched on the 2nd of June 2003 and arrived at Mars in December 2003. Since the beginning of science operations in January 2004 (Ls=330° of MY 26), Mars Express has performed continuous monitoring of conditions in the Martian atmosphere for more than 10 Martian years, and is still probing the Martian atmosphere at the time of writing. After 20 years of operations, Mars Express has provided a valuable dataset of atmospheric parameters for Mars. Exploiting the quasi-polar, non-Sun-synchronous nature of Mars Express orbit, its payload has the capability to observe the Martian atmosphere at different local times. Most importantly, the vertical structure of the Martian atmosphere can be retrieved for the whole atmospheric column, from the surface to the thermosphere, thanks to the heterogeneity and complementarity of the various instruments aboard the MEx spacecraft.

In particular, three instruments with different sensitivity in the vertical range, operating in different spectral ranges, and with different observing geometry, operate simultaneously on MEx for this purpose. The Mars Express Orbiter Radio Science (MaRS) radio occultation measurements can sound the neutral Martian atmosphere to derive vertical density, pressure, and temperature profiles as functions of height to resolutions of only a few hundred meters (Pätzold et al. 2004), allowing detailed investigation of the fine structure of the low troposphere as well as detailed characterization of the planetary boundary layer (Hinson et al. 2008; Tellmann et al. 2013). The Planetary Fourier Spectrometer (PFS, Formisano et al. 2005), a double-pendulum infrared Fourier spectrometer optimized for atmospheric studies, is well suited for retrieval of the thermal structure of the troposphere, from the surface to about 60 km, by exploiting the variation in known optical depth as a function of wavelength across the 15-μm CO₂ band (Grassi et al. 2005a; Wolkenberg et al. 2018). The relatively high spectral resolution and sensitivity of PFS, along with its wide spectral coverage, enable the simultaneous observation of numerous spectral features, ranging from optically thin lines to very optically thick Q-branches. Effectively modeling these diverse features, including absorbing and emitting Q-branches within the same spectrum observed by PFS, allows for mapping the thermal structure at various altitudes. This also facilitates the handling of atmospheric thermal inversions and strong gradients, which are well-known challenges in infrared remote sensing (Giuranna et al. 2008). The SPectroscopy for the Investigation of the Characteristics of the Atmosphere of Mars (SPICAM) experiment is an UV-IR spectrometer dedicated to nadir viewing, limb viewing, and vertical profiling by stellar and solar occultation (Bertaux et al. 2004, 2006). In its stellar occultation mode, the SPICAM UV channel can retrieve atmospheric density and temperature profiles from ~60 km up to an altitude of 160 km (Forget et al. 2009; Montmessin et al. 2017). SPICAM is, in fact, the first experiment able to provide consistent, long-term observations of the Martian mesospheric and thermospheric temperature and density profiles. These three instruments — MaRS, PFS, and SPICAM — are sensitive to different, yet complementary altitude regions. As such, they allow for a comprehensive analysis of the Martian atmospheric temperature structure. Vertical temperature profiles derived from the MaRS experiment have the highest vertical
resolution, up to 100 m, but they are especially reliable up to an altitude of ~30 km due to increasing uncertainties at higher altitudes (Tellmann et al. 2013). PFS can sound the entire troposphere with a typical uncertainty of ~1 K in the ~10–30 km altitude range, and 1–4 K at lower and higher altitudes (Wolkenberg et al. 2009; Giuranna et al. 2021). SPICAM can retrieve density and temperature profiles up to an altitude of 160 km, with an uncertainty of ~3 K around 130 km which progressively increases to larger values above and below that altitude, being larger than 10 K below 70 km (Forget et al. 2009). More details about the Mars Express mission and its scientific payload can be found in can be found in complementary papers of this collection.

In 2018, a second European spacecraft arrived at Mars: the ExoMars Trace Gas Orbiter (TGO, Vago et al. 2015). Since then, close collaboration has begun between the two missions, exploiting the synergies and complementarities of the various instruments on board the two probes. The Nadir and Occultation for MArs Discovery (NOMAD, Vandaele et al. 2018) and the Atmospheric Chemistry Suite (ACS, Korablev et al. 2018) spectrometer suites on board the ExoMars TGO have been designed to investigate the composition of Mars' atmosphere, with a particular focus on trace gases, clouds, and dust, and are intended to observe Mars' atmosphere in solar occultations, nadir and limb geometry. The solar occultation mode allows derivation of density and temperature profiles from the near-surface up to 160 km for both instruments. In addition, the ACS thermal infrared channel (TIRVIM) has been primarily dedicated to profiling temperature from the surface up to ~60 km and to monitor aerosol abundance in nadir (Korablev et al. 2018). TIRVIM has been notably used to monitor the diurnal cycle in the temperature fields in the lower atmosphere (Fan et al. 2022a; Guerlet et al. 2023). Since 2021, the Emirates Mars Mission Hope (Amiri et al. 2022) has allowed to further improve this monitoring thanks to its high-altitude orbit and its thermal spectrometer EMIRS, which has allowed to see the atmospheric thermal tides throughout the diurnal cycle every 10 sols (Fan et al. 2022b), thus extending the investigation on this topic initiated by Mars Express.

## 2 Temperature Structure of the Martian Atmosphere

Mars is a relatively small planet, yet its surface hosts enormous structures that would dwarf even the majesty of Mount Everest. A dominant feature of the planet's topography is the difference in elevation (~5 km) between the plains of the Northern Hemisphere and the lands of the Southern Hemisphere. The Tharsis region is a vast volcanic zone that interrupts the dichotomy between the northern and southern hemispheres. The largest volcanoes in the Tharsis region are 4 shield volcanoes named Ascraeus Mons, Pavonis Mons, Arsia Mons, and Olympus Mons. The latter is the tallest volcano in the entire Solar System. Its peak is about 22 km above the surrounding plain and 27 km above the average surface of Mars, with a base diameter of about 550 km. Near the equator, for about 4000 km eastwards from the center of the Tharsis region, extends the Valles Marineris, a crack in the crust of Mars about a fifth of the Martian circumference long, 600 km wide and about 7 km deep. In the southern hemisphere, with its 2300 km diameter and 9 km depth, we find the gigantic impact crater Hellas Basin (45°S, 70°E), whose bottom represents the area with lower elevation than the entire planet.

These structures are responsible for peculiar phenomena that affect the atmospheric circulation and dynamics on different scales, from local to regional and global, some of which have been observed or studied in detail for the first time by Mars Express. In this section, we highlight MEx achievements and discoveries related to the temperature structure of the
MY 26 | MY 27 | MY 28 | MY 29 | MY 30 | MY 31 | MY 32 | MY 33 | MY 34 | MY 35

<figure><img src="image_4.png" /><figcaption>**Fig. 2** Dataset of atmospheric parameters (zonal means) retrieved by PFS for 10 different MYs, including two global dust storms occurred in MY 28 and 34. Note that some biases could be introduced due to slightly different local time distributions within the averaged bins, which could be responsible for some of the differences observed in the various Martian years, especially for the ice opacity in the aphelion cloud belt (Wolff et al. 2019; Giuranna et al. 2021). Dust opacity is normalized to a reference pressure of 6.1 mbar. Reproduced with permission from Giuranna et al. (2021), copyright by Elsevier</figcaption></figure>

Martian atmosphere, grouped according to their horizontal extent. Circulation is active at all meteorological scales, which are formally divided into large-scale (> 100 s kilometers), mesoscale (> 100 s meters), and microscale (≤ 100 s meters). Here we will refer to the terms global, regional, and local in a less formal way, depending on whether the phenomenon affects the whole planet circulation (global), a particular region of the planet (regional), or just the atmosphere above a specific area on the surface (local).

## 2.1 Global Scale

The Planetary Fourier Spectrometer provided a global, long-term (almost 20 years) monitoring of the 3D temperature field in the Martian lower atmosphere, as well as of the suspended dust and ice opacity. This is one of the most extended datasets of atmospheric parameters available so far for Mars (Fig. 2). The current weather and climate of Mars has been recently studied in detail by Giuranna et al. (2021), after processing more than 2,500,000 spectra collected in more than six Martian years, covering a wide range of seasons, latitude, longitude, and local time. With unprecedented spatial and temporal coverage and details revealed, this
dataset offers new challenges to the Martian global circulation models and, at the same time, a new reference for the Martian scientific community.

Repeatable patterns are observed every year for each of the retrieved quantities, as also reported by previous instruments (see, e.g., Haberle et al. 2017, and references therein). The northern spring and summer (Ls=0–180°, also referred to as the aphelion season since aphelion (Ls=71°) occurs near the northern summer solstice (Ls=90°)) are cool, relatively dust free, and with extensive water ice clouds occurring every year (Fig. 2). On the contrary, the southern spring and summer (Ls=180–360°) are warm and dusty, relatively ice free (except for the polar hoods), and with regular occurrence of regional dust storms. The reduced insolation during the aphelion season, which is a consequence of the increased distance of Mars to the Sun due to the eccentricity of its orbit, is a primary driver of the overall cooler seasonal climate. For a detailed description of the MEx achievements and discoveries in the field of aerosol and cloud research, the reader is referred to Määttänen et al. (2024, this collection).

On the global scale, the atmospheric circulation on Mars is characterized by interhemispheric meridional Hadley cells (Wilson 1997; Forget et al. 1999). The mean meridional circulation has been investigated in detail by using zonal averages of atmospheric temperatures, as often applied in model studies, retrieved by PFS (Giuranna et al. 2021). They repeat very similarly on Mars each year, although inter-annual variations ranging from several up to tens of degrees can be observed (Fig. 3). Such variations are often, although not exclusively, related to the overall dust content in the Martian atmosphere in the different years. A regional increase of dust opacity in MY 31 is observed in two different areas: poleward of ~70°S and between 30 and 40°S latitude (black arrows in Fig. 3h). In these regions, the correlation between dust opacity and atmospheric temperature is particularly evident, with a temperature increase of ~20 K or more around ~1 mbar in correspondence with a dust opacity increase of ~0.2 (black arrows in Fig. 3h).

The Hadley cell plays a dominant role in the meridional mean circulation. Its maximum intensity occurs during the solstice seasons (around Ls=90° and 270°; central panels in Fig. 3). The ascending branch of the Hadley circulation starts in the summer tropics and transports heat to winter mid-latitudes in the descending branch. However, the polar warming, caused by the adiabatic heating of air in the descending branch of the Hadley cell, is much stronger in the northern hemisphere (during the southern summer, Fig. 3b, e, h) than in the southern one (during the northern summer, Fig. 3k, n, q). Being close to the time of perihelion, (Ls=251°), the Southern-hemisphere summer temperatures (Fig. 3b, e, h) are much larger than the Northern-hemisphere summer ones (Fig. 3k, n, q). As a result, the southern summer Hadley cell is twice more intense than the northern summer counterpart (Richardson and Wilson 2002). However, in northern summer the Hadley cell descends at ~60°S, that is about 20° of latitude more equatorward than in southern summer, allowing the southern polar vortex to be much broader than that observed in northern winter (central panels of Fig. 3).

Radio Occultation measurements usually cover small areas, which complicates the determination of large-scale or global scale atmospheric structures. Nevertheless, a sequence of several measurements in a row allows the determination of latitudinal or longitudinal large-scale structures under favorable measurement conditions. Fig. 4 (left panel) shows the latitudinal temperature structure retrieved from a sequence of MaRS measurements in the northern hemisphere during late autumn and early winter (MY 27, Ls = 253 – 274°). The temperature field shows a strong temperature inversion in the mid-latitudes associated with a strong adiabatic heating in the subsiding branch of the Hadley circulation, leading to strong warming in the upper winter polar atmosphere and strong cooling below. The thermal structure is further influenced by (breaking) gravity waves, planetary waves and thermal tides in
<figure><img src="image_5.png" /><figcaption>**Fig. 3** Daytime (8 AM ≤ LT ≤ 4 PM) zonal-mean atmospheric temperatures for bins of 60° Ls and for different Martian years, as indicated, derived from the PFS observations. The logarithmic vertical scale is used both for pressure and dust opacity (black solid line). Reproduced with permission from Giuranna et al. (2021), copyright by Elsevier</figcaption></figure>

the middle atmosphere (e.g. Barnes 1990). These temperature measurements can be used
to determine the correlated zonal wind field using the gradient wind equation (e.g. Holton
2004). Fig. 4 (right panel) shows the resulting wind field deduced from the thermal field.
A Strong zonal jet can be seen with zonal wind velocities around 170 m/s at an altitude of
around 30 km (~15 Pa).

Polar heating and cooling is driven by meridional circulation, which can also be deduced
from radio occultation measurements assuming geostrophic balance (e.g. Holton 2004). One
unique feature of radio occultation measurements is the determination of the geopotential
height on surfaces of constant pressure. Sequences of these measurements along the same
<figure><img src="image_6.png" /></figure>

**Fig. 4** left: Latitudinal temperature structure derived from MaRS radio occultation measurements (from Tellmann et al. 2013, Fig. 4). This sequence of measurements started at a latitude of 40° in late northern autumn (Ls = 253°) and moved continuously to the higher northern latitudes (lat = 75°N) close to the northern winter solstice (Ls = 274°). The local time varies from ~10:30 (latitude 40°N) to ~15:00 (latitude 75°N). A strong temperature inversion can be seen. This temperature inversion is correlated with a strong zonal jet deduced from the gradient wind equation (right panel, from Tellmann et al. 2013, Fig. 6). Right: Zonal winds derived from temperature structures in the left panel using the gradient wind equation and the latitudinal pressure gradient. Reproduced with permission from Tellmann et al. (2013), copyright by AGU

**Fig. 5** Meridional winds (positive northward) retrieved under the assumption of geostrophic balance from the geopotential height profiles located at a small latitude band along various longitudes. Solid lines indicate the zero line. Reproduced with permission from Tellmann et al. (2013), copyright by AGU

<figure><img src="image_7.png" /></figure>

latitude can be used to determine meridional wind speeds, also known as geostrophic winds. Fig. 5 shows a sequence of measurements aligned along the same latitude at various longitudes. These meridional wind structures indicate an effective poleward heat transport along certain longitudes, as well as an equatorial transport of cold air at longitudes of ~90° and 350° east. The measurements were retrieved in MY 30 at latitudes between 49°S and 69°S. The measurements are located around the southern winter solstice (Ls = 61-110°) in the early afternoon (LT ~13:00 - 14:00).

SPICAM has been able to probe the Martian atmospheric thermal structure from 60 to 130 km by making use of stellar occultation in the UV range. This kind of measurement consists in targeting a UV-bright star in the sky and staring at it through the Martian atmosphere while the star rises or sets behind the horizon. The UV range covered by SPICAM (118 to 320 nm; see Montmessin et al. 2017) covers a broad and deep CO₂ absorption band that can be identified by SPICAM up to 130-150 km (number dependent on season and region), which in turns allows density and temperature to be profiled. However, the dusty Martian atmosphere is too opaque below 50 km to allow CO₂ to be identified in the tro-
<figure><img src="image_8.png" /><figcaption>**Fig. 6** The latitudinal distribution of the occultations as a function of $L_s$. The black solid line shows the latitude of the subsolar point. Shaded areas illustrate the extension of the polar nights. Reproduced with permission from Forget et al. (2009), copyright by AGU</figcaption></figure>

posphere. Therefore, SPICAM temperature profiles are restricted to the middle and upper atmosphere, which had remained poorly explored before Mars Express. As indicated by Forget et al. (2009), “the few observations available (before SPICAM) have revealed a rich and complex dynamics combining various kinds of waves probably interacting with the mean circulation”. These authors performed an analysis of the first 616 SPICAM density and temperature profiles that covered MY27. Stellar occultation offers more opportunities than solar occultation due to the multiplicity of potential targets; yet the necessity to explore mostly the nightside (to avoid solar scattered light) combined with satellite operation constraints (the pointing being performed by the spacecraft) made the coverage uneven in space and time (Fig. 6). A pole-to-pole excursion could only be accomplished within a Martian season (90° of $L_s$), that is, the typical duration over which the “global” Martian atmosphere could be appreciated. Forget et al. (2009) identified an unexpected increase in atmospheric density at 60 km at $L_s$ 130° when dust storm activity is usually low. While a detailed exploration of the diurnal cycle in the mid-to-upper atmosphere was not possible, some day-to-night temperature differences were perceptible in the data, consistent with propagating thermal tides and solar heating variations. A striking characteristic of the temperature near the mesopause (supposed to be the coldest atmospheric region) was the numerous profiles exhibiting temperatures colder than CO₂ frost point (6% of the cases), which followed up the discovery of mesospheric CO₂ ice clouds in the same altitude range and discussed hereafter. This altitude range also revealed that atmospheric general circulation model predictions were too warm by approximately 30 K, which appeared to be caused by an underestimation of the atomic O₂ concentration which controls the CO₂ infrared cooling.

An expansion of this work was presented in Montmessin et al. (2017). Several factors contribute to the large seasonal fluctuations in atmospheric densities, as analyzed by Forget et al. (2009), including variations in dust content in the lower atmosphere, which affects temperature, and thus atmospheric scale height, below 50 km. According to Fig. 7, there is a direct correlation between the upper air density and the lower air dust loading, since a linear function of the visible dust opacity measured by the camera on the Opportunity rover (Lemmon et al. 2004) can match the variations of the observed density at 100 km.

The response of the upper atmosphere of Mars to lower atmospheric dust storms has also been studied using SPICAM data and other datasets by Withers and Pratt (2013). The SPICAM-model disagreement found near the mesopause was almost completely solved by adding an improved parameterization to the model that was specifically developed to predict colder mesopause temperatures.

In some specific conditions, the temperature in the atmosphere of Mars can be low enough that CO₂ condenses into ice clouds. In most cases, those conditions are reached in
<figure><img src="image_9.png" /><figcaption>**Fig. 7** (Black dots) CO₂ density measured by stellar occultation at 100 km above the Mars zero datum (areoid) between January 2004 (MY26, Ls = 333°, here shown at Ls = -27°) and February 2011 (MY29, Ls = 238°, here shown at Ls = 958°). Only the data obtained below 50° latitude are shown. (Green curve) A synthetic density ρ=(5τ -2) × 10⁻⁷ directly derived from the visible aerosol opacity τ measured in Gale crater (Lemmon et al. 2004). Reproduced with permission from Montmessin et al. (2017), copyright by Elsevier</figcaption></figure>

the troposphere in the polar night, in the mesosphere at the equator before and after aphelion, and around perihelion in the southern mid-latitudes. Tropospheric clouds were reported in the winter polar night with the Mars Orbiter Laser Altimeter (MOLA) onboard Mars Global Surveyor (MGS) (Pettengill & Ford 2000) at locations corresponding to the low brightness temperatures measured from the Thermal Emission Spectrometer (TES) and thus likely candidates for CO₂ ice clouds (Ivanov & Muhleman 2001). However, the spectral signature of CO₂ ice clouds was first detected in PFS limb-viewing mesospheric observations at 80-85 km, 138° Ls, and 11:30 local solar time (Formisano et al. 2006). A spectral signature of CO₂ ice cloud in the polar regions was also identified from MCS/MRO measurements (Hayne et al. 2012). The presence of CO₂ ice clouds was inferred by other instruments on MEx such as the UV channel of SPICAM, which detected very high clouds in southern winter subtropical latitudes in the nightside, corresponding to temperatures falling below the CO₂ condensation limit (Forget et al. 2009; Montmessin et al. 2006). Figure 8 shows some examples of temperature profiles retrieved from the SPICAM-UV instrument with some values lower than the limit for CO₂ condensation. All those temperature minima are coupled with a detached layer of aerosols. It was noticed that the CO₂ cloud could stretch below the 'cold pocket' as CO₂ ice falls and does not directly sublimate entirely when reaching warmer regions.

The CO₂ ice spectral signature at 4.26 µm was also observed by the OMEGA mapping spectrometer (Observatoire pour la Minéralogie, l'Eau, les Glaces et l'Activité; Bibring et al. 2004) in the equatorial region (Montmessin et al. 2007), before and after the northern summer solstice (Ls 45° and 135°). In addition to the temperature vertical profiles from SPICAM, those from the Thermal Emission Spectrometer (TES) and Radio Science indicate the presence of very cold pockets where CO₂ can condense into ice (Colaprete et al. 2008). The temperature retrieved from the Mars Climate Sounder (MCS) on MRO was inspected in Sefton-Nash et al. (2013) and they suggested that a limit of 20 K or less above the CO₂ frost point might serve as a reasonable indicator of possible CO₂ ice formation. The temperature needed for nucleation should be lower than the limit for CO₂ condensation for initiating heterogeneous nucleation while the temperature required for homogenous nucleation would be much lower and has never been reported in Mars atmosphere (González-Galindo et al. 2011). The presence of CO₂ ice clouds was further inferred by different instruments such as TES and the Mars Orbiter Camera (Clancy et al. 2007), OMEGA and HRSC (Montmessin et al. 2007; Määttänen et al. 2010; Scholten et al. 2010), THEMIS on Mars Odyssey (Inada
<figure><img src="image_10.png" /><figcaption>**Fig. 8** Four temperature profiles retrieved from SPICAM-UV with temperature values below the limit for CO₂ condensation (dashed). The horizontal dash-dot lines correspond to the top of some opacity due to the presence of aerosols along the line of sight. The presence of the aerosol layer might seem lower than the minimum in temperature, depending on its location along the slant path. The temperature minimum can extend far away from the tangent point. Reproduced with permission from Montmessin et al. (2006), copyright by Elsevier</figcaption></figure>

et al. 2007; McConnochie et al. 2010), OMEGA and CRISM (Vincendon et al. 2011), IUVS on MAVEN (Jiang et al. 2019; Stevens et al. 2017), PFS and OMEGA (Aoki et al. 2018), CRISM, MCS, and MARCI (Clancy et al. 2019) and, more recently, at Mars terminator by the NOMAD-SO channel (Liuzzi et al. 2021). The CO₂ ice clouds and the ‘cold pockets’ seem to be limited to specific seasons and locations, around and between Meridiani and Valles Marineris (Clancy et al. 2019; Määttänen et al. 2010; Vincendon et al. 2011), and east of Hellas Basin at southern mid-latitudes (Liuzzi et al. 2021). Most of these times and locations are either in the equatorial region (20°S to 20°N) around northern summer or at mid-latitudes in the southern hemisphere at solar longitudes of 200° to 300° and the cold pockets seem to be related to local mechanisms such as gravity waves (Scholten et al. 2010; Spiga et al. 2012; Slipski et al. 2022) in combination with thermal tides (González-Galindo et al. 2011; Jiang et al. 2019; Listowski et al. 2014). Both times and locations of the cold
<figure><img src="image_11.png" /><figcaption>**Fig. 9** (a) Impact of MY 28 dust storm on daytime (8 AM ≤ LT ≤ 4 PM) atmospheric temperature observed by PFS (from Giuranna et al. 2021). (b) Net heating and cooling rates retrieved from PFS observations. Dates and retrieved dust opacities are labeled in the Figure. Reproduced with permission from Wolkenberg et al. (2018), copyright by Elsevier</figcaption></figure>

pockets are well reproduced by GCMs (González-Galindo et al. 2011) but the temperatures simulated were too high to infer CO₂ condensation.

## 2.1.1 Global Dust Events

Dust is a highly variable component of the Martian atmosphere, which exerts a strong influence on climate and meteorology. Dust storms on Mars can be found in both the northern and southern hemispheres, and their occurrence tends to be enhanced during southern spring and summer, when Mars orbits closer to the Sun. Dust storms are typically regional in extent, yet can aggregate and develop into planet-encircling events (hereafter referred to as Global Dust Event; GDE), one of the most spectacular phenomena in our solar system. By absorbing the solar radiation, dust can efficiently warm the lower atmosphere by diabatic heating and produce horizontal and vertical expansion of the meridional circulation, which can affect thermal tides and even the global circulation (Zurek et al. 1992; Haberle et al. 1982).

GDEs originate near perihelion from a combination of multiple local and regional dust-lifting events (Smith et al. 2002; Haberle et al. 2017) and have been observed in detail by several orbiters and landers. The decay of the MY 9 GDE was observed by Mariner 9 in 1971 (Conrath et al. 1973; Leovy et al. 1971; Pang and Hord 1973) while the Viking orbiters and landers observed the GDE occurred in MY 12 (Briggs et al. 1979; Ryan and Sharman 1981). The next GDE occurred in MY 25 and was observed in detail by MGS (Smith 2004). The most recent Global Dust Events (GDEs) on Mars occurred during Mars Years (MYs) 28 and 34. These events have been extensively observed by several orbiters and rovers, including the Mars Reconnaissance Orbiter (MRO), Mars Odyssey (MO), Mars Express (MEX), and the Mars Exploration Rovers. As a result, they are the most thoroughly studied storms to date (Fedorova et al. 2018; Guzewich et al. 2017, 2020; Heavens et al. 2019; Kass et al. 2019; Kleinböhl et al. 2020; Lemmon et al. 2015, 2019; Smith 2009, 2019; Wang and Richardson 2015; Wolkenberg et al. 2018, 2020).

During MEX operations, two global dust storms have developed on Mars (Fig. 2), namely in MY28 and MY34, allowing the effects of suspended dust on atmospheric temperatures to be studied at those times. Giuranna et al. (2021) compared the atmospheric temperatures recorded by PFS during the GDE of MY 28 with those observed over the same seasonal range (240–300° Ls) for a “typical” non-GDE Martian year (Fig. 9a). During the GDE, the atmosphere was strongly heated. For altitudes corresponding to pressure levels higher than ~1 mbar, the mean atmospheric temperature was found to be 28 K warmer than during a
non-GDE year. However, PFS also observed a net cooling of the lower atmospheric layers of up to 26 K, on a zonally averaged sense (Fig. 9a). The important result here is that the infrared radiative cooling rates due to dust cannot be overlooked in the lower altitudes within a dusty atmosphere; since cooling in the infrared regions is comparable to heating due to absorption of the incident solar radiation by dust, and may actually be thought that it dominates the effect on temperature since it affects the denser layers of the atmosphere (Fig. 9b).

The MY 34 GDE, which peaked around Ls 185-210°, was in parallel observed by TGO instruments both in nadir and solar occultation. The temperature field in nadir viewing was measured by the TIRVIM channel of ACS (Vlasov et al. 2022), an FTS resembling the long-wavelength PFS channel, and covering the spectral range from 600 to 1300 cm⁻¹ with a resolution of 1.2 cm⁻¹. The TIRVIM-derived temperature profiles range from the surface up to 55-60 km; the column dust optical depth was also retrieved from the silicate absorption band. TIRVIM observed how the global storm caused an asymmetric heating of the atmosphere, mostly in the southern hemisphere, and changed the amplitude of the diurnal temperature contrast both in the atmosphere and at the surface.

## 2.2 Regional Scale

Some of the largest topographical features present on the Martian surface, such as Hellas basin, Olympus Mons and Valles Marineris, are often responsible for peculiar phenomena that affect the atmospheric circulation and dynamics on a regional scale.

**Mars Polar Cap "mystery" solved** The residual south polar cap is approximately 400 km across, and the cap center is displaced by about 3° from the geographic pole. Mars eccentricity or local topographic irregularities cannot straightforwardly account for this asymmetry, nor the energy balance models can reproduce it. In fact, this has been a major issue in Mars science for many years. The condensing CO₂ south polar cap of Mars and the mechanisms of the CO₂ ice accumulation have been studied through the analysis of spectra acquired by PFS after the first two years of MEx mission (Giuranna et al. 2008). PFS found that, in the Ls = 50°–70° period, the south pole of Mars is characterized by two distinct regional climates: CO₂ snow falls are allowed exclusively in the western hemisphere, where the atmospheric temperatures go below the condensation temperature of the gaseous CO₂ at given altitudes (Fig. 10). In the eastern hemisphere the whole atmospheric column is several degrees above the saturation temperature (Fig. 10) and the cap consists mainly of CO₂ frost deposits, as a consequence of direct vapor deposition.

The residual cap lies in the cold region, where CO₂ snowfalls occur. The longitudinal thermal asymmetries seen between the western and eastern hemispheres in southern high latitudes are the result of a planetary wave that is excited primarily by the middle-latitude surface topography, namely the Hellas and Argyre basins (Colaprete et al. 2005). This wavenumber-one planetary wave establishes a high-pressure zone over much of the eastern hemisphere, preventing CO₂ condensation in the atmosphere. Unward motions occur in the western hemisphere, establishing a cold, low-pressure zone. This makes the CO₂ precipitation very likely to occur in the western hemisphere. The two climates established by the topographically forced circulation produce differences in the nature and amount of the surface ice deposits during the mid-late fall. Thus, topography, through its influence on the circulation, is ultimately responsible for the nature and location of the residual South polar cap. Frost grain size is determined by the style of frost deposition. Smaller grains form when there is atmospheric precipitation to the surface, and larger grains form when frost
<figure><img src="image_12.png" /><figcaption>**Fig. 10** (left) $L_s = 50°–70°$. Longitude cross-section at –85° latitude showing the vertical profile of $T – T_{sat}$ (the CO₂ condensation temperature). At these $L_s$, the wavenumber-one planetary wave is already well-defined and so are the two distinct climates on either side of the pole. (right) CO₂ snowfalls are allowed exclusively in the western hemisphere. The coldest region [blue arrow in panel a] is located right around the center of the south residual cap (blue dotted line). Reproduced with permission from Giuranna et al. (2008), copyright by Elsevier</figcaption></figure>

is directly deposited to the surface. Smaller grains result in higher albedo and vice versa (Calvin and Martin 1994; Warren et al. 1990). Moreover, the higher the albedo, the lower the sunlight absorbed and, consequently, the sublimation rate (James et al. 1992; Thomas et al. 1992). Therefore, the different amounts of CO₂ ice deposited in the two hemispheres during the fall and winter seasons, together with the different sublimation rates due to different grain sizes and, in turn, different albedo between CO₂ (fresh) snow and frost, explain the asymmetry observed at the end of the southern summer: the dry ice sublimes entirely in the eastern hemisphere, while in the western hemisphere it survives all year long.

**The fountain of Olympus** With its 25 km height, Olympus Mons has a strong impact on the atmospheric circulation at a regional scale, especially if we consider that the scale height for a temperature variation on Mars is about 10-11 km. The atmosphere over great volcanoes on Mars has been largely studied from measurements of several instruments (Zasova et al. 2001, 2006; Benson et al. 2003, 2006; Wang and Ingersoll 2002; Pearl et al. 2001; Grassi et al. 2005b) as well as by mesoscale models (Rafkin et al. 2002; Michaels et al. 2006; Moudden and McConnell 2005; Spiga and Forget 2009). The atmospheric temperature profiles retrieved by PFS have been used to study the thermal structure and the circulation of the Martian atmosphere in an area 50° × 40° (lat × lon) around Olympus Mons (Wolkenberg et al. 2010). Three striking features were observed: two warm regions North and South of the Volcano at 8°N and 26°N (Fig. 11, black dotted lines), which are referred to as thermal bumps or “hot ring” (see Fig. 12); two regions almost vertically isothermal which form a collar of air that is colder than nearby at low altitudes, and warmer than nearby at high altitudes (red dotted lines in Fig. 11); and one region on top of the volcano with a high temperature just above it and rapidly decreasing with altitude, where the coldest temperature of the entire region is also observed (blue dotted line in Fig. 11).

To understand the meaning of the three features mentioned above, the corresponding atmospheric temperature profiles have been analyzed in detail. The air parcel on top of Olympus Mons is thermally buoyant. The potential temperature there is decreasing with altitude (negative slope for at least 5–7 km) which is an indication of an unstable atmosphere.
<figure><img src="image_13.png" /></figure>

**Fig. 11** Example of atmospheric thermal fields around Olympus Mons observed by PFS between 9°S and 40°N latitude. Isothermal contours in the atmosphere are shown (every 10 K) as a function of altitude (in km, vertical scale) and time (in seconds: upper scale) or latitude (in degrees: bottom scale). The bottom-line profile is the topography obtained from the MOLA altimetry for the location of PFS observations. Reproduced with permission from Wolkenberg et al. (2010), copyright by Elsevier

**Fig. 12** Atmospheric temperatures around Olympus Mons for various altitudes. 2-5 km: the “Hot Ring”; 10 km: the “Cold Collar”; 20 km: the “Hot Spot” on top of the Volcano. Reproduced with permission from Wolkenberg et al. (2010), copyright by Elsevier

<figure><img src="image_14.png" /></figure>

The temperature profiles inside the two warm regions follow well the adiabatic lapse rate in the last 10 km. Finally, the profile relative to the cold collar around the volcano, inside the hot ring, shows less than half the vertical temperature variation of the other profiles, justifying the pseudo 'isothermal' impression obtained from Fig. 11. The final interpretation is that the hot spot on top of the volcano acts as the motor for a peculiar regional circulation. Due to the solar heating, the local air parcels on top of the volcano reach a temperature so high that they become unstable and move upward, with strong adiabatic expansion. A cold collar then is generated at the base of the volcano: air parcels are sucked upward and adiabatically cooled. At some level, the air parcels are pushed sideways, resulting in a quasi-isothermal vertical profile. Finally, the air parcels fall downward and are compressed adiabatically. They
<figure><img src="image_15.png" /><figcaption>**Fig. 13** (a) Morning fog in Valles Marineris observed by HRSC at L<sub>S</sub> = 38°. (b) PFS measured spectra inside (133), on the edge of (136) and outside (130) Valles Marineris. Spectrum 133 shows lower temperatures and the water ice band at 750–900 cm<sup>-1</sup>. (c) Vertical temperature profiles obtained from three PFS observations: south of (dotted line), inside (dot-dashed line) and north of (dashed line) Vallis Marineris. Note the thermal inversion inside Vallis Marineris. From Möhlmann et al. (2009)</figcaption></figure>

follow the adiabatic lapse rate and form a “hot ring” around Olympus Mons (Fig. 12). This
meso-scale circulation has been called “the fountain of Olympus”.

**Fog phenomena on Mars** Valles Marineris is home to another spectacular phenomenon observed by Mars Express. The High-Resolution Stereo Camera (HRSC, Jaumann et al. 2007) and PFS have documented and thermodynamically verified the existence of water ice fog in the near-surface atmosphere of Mars, and particularly inside Valles Marineris (Möhlmann et al. 2009). An example is shown in Fig. 13a, where the presence of fog is evident from the HRSC image. PFS demonstrated that the appearance of fog in Valles Marineris can be explained in terms of fog of water ice particles, due to the presence of the water ice band in the spectra corresponding to where the fog is observed in the HRSC images (Fig. 13b). Moreover, PFS temperature profiles show striking differences if observed inside, on the edges, or outside Valles Marineris: a thermal inversion is observed only inside Valles Marineris, in connection with the presence of the morning fog (Fig. 13c). This aspect favors even more the formation of the morning fog as it damps vertical mixing. Finally, numerical estimates of the water vapor pressure, and the related atmospheric water content, at the frost-point (temperature of freezing or “re-sublimation” of water vapor) by a 1-dimensional planetary boundary layer model indicate that conditions in the planetary boundary layer can indeed temporarily favor the formation of ice particles (Möhlmann et al. 2009).

**2.3 Local Scale**

**The planetary boundary layer** The interaction between the planetary surface and the lower atmosphere plays a crucial role in the lowest atmospheric layer, the planetary boundary layer
<figure><img src="image_16.png" /><figcaption>**Fig. 14** Four Profiles of temperature (a) and potential temperature (b) located at Amazonis Planitia (a, c) and above Tharsis (b, d) (MY 28, Ls~25°, LT ~15.5 h). The lower panel shows the location and the topography. Reproduced with permission from Tellmann et al. (2013), copyright by AGU</figcaption></figure>

(PBL). The solar heating of the surface leads to the efficient heating of the adjacent atmo-sphere and the formation of a convective boundary layer (CBL) which shows a distinctive diurnal variability, as shown by MaRS. The low atmospheric density especially over elev-ated terrain allows it to react especially fast to diurnal temperature changes, leading to the formation of deep CBLs over elevated terrain. Hinson et al. (2008) showed a clear corre-lation between surface elevation and the depth of the CBL. The depth of the CBL over the elevated tropical terrain reaches values of up to 9 km, while it only reaches a depth of ~3 km in the northern lowlands (Hinson et al. 2019). Fig. 14 shows typical temperature and poten-tial temperature profiles retrieved in the lowlands (Amazonis Planitia (a, c)) in comparison with profiles above Tharsis (b, d). The depth of the CBL above tropic elevated terrain can reach values of more than 9 km while the CBL is very shallow over the midlatitude lowlands (lower than 3 km depth) (Hinson et al. 2019). The temperature lapse rate in the CBL is nearly adiabatic as can be seen by the nearly static stability which is close to zero in this region (Fig. 14). Numerical Large Eddy Simulations (i.e. atmospheric model with horizontal reso-lution lower than 100 m) show that, in the tenuous CO₂ Martian near-surface environment, the daytime BL is to first order controlled by the infrared radiative heating, fairly indepen-dent of elevation, which implies a simple correlation between the BL potential temperature and the inverse pressure ("pressure effect", Spiga et al. 2010).

**Gravity waves** Gravity waves are recognized for their significant role in redistributing energy and momentum on Earth (e.g., Fritts and Alexander 2003), and atmospheric models suggest that they also exert a substantial influence on the atmospheres of other terrestrial
<figure><img src="image_17.png" /><figcaption>**Fig. 15** Temperature fluctuations from MaRS over four selected regions in northern spring, MY27 (LT: ~17 h). Upper panel: Measurements over low-lands, lower panel: measurements at the Tharsis ridge. Reproduced with permission from Tellmann et al. (2013), copyright by AGU</figcaption></figure>

planets. These atmospheric waves are triggered by various source mechanisms, with buoyancy serving as the restoring force for these small-scale waves. Key source mechanisms include air displacement over topographical obstacles, convection, wind shear instabilities, and near-surface thermal contrasts.

The high vertical resolution of MaRS profiles enables the retrieval of small-scale vertical temperature disturbances, primarily induced by gravity waves. In Fig. 15, typical temperature fluctuations measured over two distinct topographical regions are displayed. The exponential growth of amplitude with height signifies freely propagating gravity waves. These waves can ascend to a level of atmospheric instability, at which point they break or saturate. It is believed that breaking gravity waves play a crucial role in atmospheric circulation (e.g., Barnes 1990). Gravity wave activity over elevated terrain (Tharsis ridge, lower panel) is considerably more vigorous than in the lowland region (upper panel), despite the topographical dependency being lower than predicted by models (e.g., Creasey et al. 2006).

Another significant factor influencing the formation of gravity waves is the seasonal variability of the atmosphere. Strong zonal jets in the winter hemisphere can result in a pronounced increase in gravity wave activity, particularly around winter solstice.

**Effects of a local dust storm** Large-scale dust storms on Mars occur every year, especially during the “dusty season” (second half of the year). However, local dust storms may occur at
<figure><img src="image_18.png" /><figcaption>**Fig. 16** The latitude-pressure cross-sections of the temperature structure retrieved from PFS LWC observations for orbit #1201 (top) and #1212 (bottom). The positions of the temperature profiles (shown in the right panel) are marked with lines: the solid line is from the dusty orbit #1212 inside the dust storm, the dashed line from the dust-free orbit #1201 at the storm area, and the dash-dotted line is from the dusty orbit outside the dust storm. The local time of observations at the dust storm region was 13.30. Reproduced with permission from Määttänen et al. (2009), copyright by Elsevier</figcaption></figure>

any time and are particularly difficult to observe due to their small extent and short duration. One needs to observe the planet at the right place at the right time to be able to observe and study the properties of a local dust storm and its effect on the atmospheric temperature. Määttänen et al. (2009) used the observations of two Mars Express instruments, OMEGA and PFS, to distinguish and analyze a local dust storm near the equator of Mars in late northern summer. The same geographical unit was observed on two orbits only 3 sols apart, and the later orbit revealed a local dust storm, whereas in the first orbit the atmosphere was very clear (Fig. 16). The temperature profiles retrieved by PFS revealed a net cooling close to the surface and a net heating in the upper atmosphere due to the presence of suspended dust, which is consistent with previous theoretical considerations (Gierasch and Goody 1972). In the core of the local dust storm, the temperatures observed near the surface and at 10 Pa were ~10 K cooler and ~5 K warmer, respectively (blue and red arrows in Fig. 16). As a result, atmospheric dust tends to homogenize the vertical temperature profile, although this trend has a weaker amplitude in local dust storms than in regional and global dust storms.

**MEx support to NASA's Mars Landers** On May 25, 2008, NASA's Phoenix Mars Lander landed in the northern polar region of the red planet. Flight controllers at ESA's Space Operations Centre optimized the Mars Express orbit so that the spacecraft could provide good visibility of Phoenix's planned trajectory. To help characterize the descent trajectory, MEx provided critical communications back-up and PFS was used to survey the Martian atmosphere before and after descent to better characterize how the descent trajectory was affected by the atmosphere. On February 18, 2021, NASA's Mars 2020 Perseverance rover landed on the Red Planet. A large team of scientists provided support for the landing, utilizing essentially all available data sources. The majority of the data came from MCS and MARCI, both on MRO, while several other instruments, including THEMIS-MO and PFS-MEX, provided
<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /><figcaption>**Fig. 17** (Left) Image of Jezero crater with the center of the PFS observation “spots” within Jezero crater and instrument footprint. (Right) Temperature profiles obtained from PFS measurements. These measurements were obtained on 2021 February 14, 4 days prior to EDL. Reproduced with permission from Mischna et al. (2022), copyright by the author(s)</figcaption></figure>

relevant data on the thermal structure and opacity of the atmosphere around Jezero Crater. Both ESA's Mars orbiters – TGO and MEx – supported the landing. TGO relayed important data from Perseverance to Earth as soon as four hours after landing, while Mars Express monitored the local conditions at the landing site, Jezero crater. Having up-to-date knowledge of the temperature, pressure, dust and ice conditions in the atmosphere is crucial for understanding its density and predicting and analyzing the trajectory of the rover's descent to the surface of the red planet. PFS was able to provide useful temperature profiles on an intermittent basis in the weeks leading up to EDL and a daily basis during the final week of approach, predominantly in its “spot-tracking” mode (Fig. 17), where the instrument focuses on a single location on the surface as it passes overhead, providing multiple profiles as a function of time (Mischna et al. 2022). The information collected by PFS was passed on to the Perseverance Entry, Descent and Landing (EDL) team at NASA and included in their daily report during the two weeks leading up to the landing.

## 3 Discussion and Conclusions

Mars Express, in particular with its instruments PFS, MaRS and SPICAM, has provided a unique dataset on the thermal structure of the Martian atmosphere, complemented later by the following Trace Gas Orbiter. The lower atmosphere (up to 40-50 km) has been profiled globally over 20-years by PFS at various local times, and locally by MaRS radiooccultations at high vertical resolution. The thermal structure of the upper atmosphere (up to 130 km) has been revealed for the first time by SPICAM UV stellar occultations, later completed by TGO NIR solar occultations (up to 180 km).
These data can and have been used in two ways: 1) as a reference dataset on the climate of Mars, available to make reference climatologies and validate climate models. 2) to reveal and explain specific processes (waves, convection, hazes and clouds, etc.) as shown in many examples in this paper. There is no doubt that much more scientific discoveries are awaiting in the Mars Express dataset, in particular by using the observations in combination with other datasets or numerical models.

After decades of observations and analysis with various numerical models, it is tempting to say that the thermal structure of the Martian atmosphere is understood, which is true to first order. Global Climate Model can predict well most of the observed temperatures within a few Kelvins on the basis of universal physical equations, when the variations in dust loading are properly observed. Models may fail to accurately predict the observed temperature when not driven by observations of atmospheric dust. However, the remaining discrepancies are meaningful. They certainly reveal profound knowledge gaps in the climate processes at work. These may include the behavior of waves and wave-mean flow interactions, the microphysics of clouds, the radiative properties of dust (are dust particles always in radiative equilibrium with the air?) and its transport, the behavior of the boundary layer in the extreme stable and unstable Martian conditions, etc. As mentioned in this paper, temperature measurements are also used to estimate the winds using the gradient wind equation or via numerical models and/or data assimilation. However, because of the observation uncertainties and the complex control of the winds by waves-mean flow interaction, it is likely that direct wind measurements will be required to validate our understanding of the Martian atmospheric circulation that has mostly been constrained via temperature measurements. Meanwhile, more should be done with the unique datasets obtained by Mars Express and the TGO missions.

**Acknowledgements** The authors thank the International Space Science Institute (ISSI) in Bern, Switzerland for the support.

**Author Contribution** MG drafted the first version of the review paper. All authors participated in the writing of the content in line with their expertise in the field, and their connection to the missions and instruments described. All authors commented on previous versions of the manuscript. All authors read and approved the final manuscript.

**Funding Information** Open access funding provided by Istituto Nazionale di Astrofisica within the CRUI-CARE Agreement. The PFS experiment was built at the Institute for Space Astrophysics and Planetology (IAPS) of the National Institute for Astrophysics (INAF) and is currently funded by the Italian Space Agency (ASI; agreement number 2018-2-HH.0). Parts of the studies presented in this paper received funding from the European Union's Horizon 2020 Programme (H2020-COMPET-08-2014) under grant agreement UPWARDS-633127. The MaRS experiment was funded by the Deutsches Zentrum für Luft- und Raumfahrt (DLR) under contract 50QM1802 and contracts with NASA. The MaRS experiment has benefited greatly from the continued support of the ESA ESTRACK ground station crews and the NASA Deep Space Network. SPICAM was developed and is operated in LATMOS (CNRS/Univ. Paris-Saclay, France), has been funded by the Centre National d'Études Spatiales (CNES) in France; its IR channel was conceived and developed at the Space Research Institute (IKI) in Russia. The NOMAD experiment is led by the Royal Belgian Institute for Space Aeronomy (BIRA-IASB), with Co-PI teams in the United Kingdom (Open University), Spain (IAA-CSIC) and Italy (INAF-IAPS). The ACS experiment is led by IKI in Russia, supported by the LATMOS Co-PI team and funded by Roscosmos. Some part of this project has received funding from the European Union's Horizon 2020 research and innovation programme under grant agreement No 101004052.

**Declarations**

**Competing Interests** The authors have no relevant financial or non-financial interests to disclose. The authors have no competing interests to declare that are relevant to the content of this article.
**Open Access** This article is licensed under a Creative Commons Attribution 4.0 International License, which permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons licence, and indicate if changes were made. The images or other third party material in this article are included in the article's Creative Commons licence, unless indicated otherwise in a credit line to the material. If material is not included in the article's Creative Commons licence and your intended use is not permitted by statutory regulation or exceeds the permitted use, you will need to obtain permission directly from the copyright holder. To view a copy of this licence, visit http://creativecommons.org/licenses/by/4.0/.

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