Dust climatology from NOMAD UVIS channel
Z. Flimon1, J. Erwin1, A.C. Vandaele1, L. Neary1, A. Piccialli1, L. Trompet1, Y. Willame1, F. Daerden 1, S.
Bauduin2, I. R. Thomas1, B. Ristic1, J. Mason3, C. Depiesse1, M. R. Patel3, G. Bellucci4, J.-J. LopezMoreno5
1

Royal Belgian Institute for Space Aeronomy, BIRA-IASB
Université libre de Bruxelles (ULB), Spectroscopy, Quantum Chemistry and Atmospheric Remote Sensing
(SQUARES), Brussels, Belgium
3
School of Physical Sciences, The Open University, Milton Keynes, UK
4
Instituto de Astrofisica e Planetologia Spaziali, INAF, Rome, Italy
5
Instituto de Astrofìsica de Andalucía, Consejo Superior de Investigaciones Científicas (CSIC), Granada, Spain
2

Introduction:
Aerosols present in the atmosphere of Mars have
a major effect on it. They are mainly composed of
dust, water ice or CO2 ice. Dust is confined to lower
altitudes during the aphelion season and can reach
higher altitudes during the perihelion, especially during dust storms that frequently arise on Mars. These
storms can sometime grow up to cover the entire
planet and are then called a global dust storm.
Notable observed water ice cloud features include
polar hood and the Aphelion cloud belt (Clancy et
al., 1996); CO2 ice clouds can be observed seasonally at high altitudes (Montmessin et al., 2006),(Liuzzi
et al., 2021). Where water ice clouds can be found
mainly during perihelion and CO2 cloud during aphelion. Dust can be found all the year round at low altitude and represent the main component of Martian
aerosols.
As explained before dust can be transported to
high altitudes during perihelion and dust storms.
Dust in the atmosphere can have different effects, for
example during dust storms it can be responsible for
hydrogen loss (Holmes et al., 2021). Also, one of the
most important effects of dust is the local atmospheric warming: the sun heats the dust which then heat up
the local atmosphere. Knowing the vertical profile of
dust is important to know the temperature profile.
Aerosols climatology has been studied by previous missions in solar occultation with SPICAM on
board Mars Express (Määttänen et al., 2013),
(Fedorova et al., 2014) or with limb measurement
with CRISM. (Clancy et al., 2019). In this work we
will focus our analysis on the UVIS channel of
NOMAD.
The NOMAD (“Nadir and Occultation for MArs
Discovery”) spectrometer suite on board the
ExoMars Trace Gas Orbiter (TGO) is composed of
three spectrometers, two in IR (LNO and SO) and
one in UV-visible (UVIS).
The UVIS spectral range extends from 200nm to
650nm with a spectral resolution about 1.5nm. UVIS
can operate in nadir and occultation modes. In this
work, we use observations taken in occultation mode
to investigate the vertical distribution of aerosols.
The altitude resolution of UVIS is below 300m and
the SNR greater than 1000 for wavelength between

230-450nm, and generally larger than 500 for wavelength between 450-650nm (Vandaele et al., 2015).
Using the UVIS channel, this work will permit to
have a great precision when studying the vertical
profile of aerosols and also spectral features.
We analyzed data from 2018 to 2021 covering 2
Martian years (see figure 1).

Figure 1 : Repartition of UVIS solar occultation
between Martian years 34 to 36
We cover practically all the latitudes, this will
give us the data necessary to build a complete climatology of aerosols in the Martian atmosphere.

Method:
To compute aerosol’s extinction, from the transmittances of UVIS. We use first the ASIMUT code
(Vandaele, et al., 2006.) to make a fit on ozone and
Rayleigh scattering and then subtract them from the
original transmittances. In the result should remain
only the background of the spectra. Extinction can be
computed from the transmittance after subtraction of
ASIMUT’s fit using the formula from (Wilquet et al.,
2012):
)
(
With τ the optical depth , T the transmittance, I the
solar irradiance attenuated through the atmosphere
and I0 the reference irradiance of the solar spectrum

outside the atmopshere. β represents the extinction
and N the number of layer above the current layer n.
λ represents the wavelength and dZ represents the
pathlength of light through the atmosphere to the
point and i represent the upper layer.
To confirm the validity of our observations, only
the data with a SNR > 8 and a minimum of 100
points in the spectra are kept.
The extinction is fitted using the refractive index for
Mars dust from (Wolff et al., 2009). The extinction
efficiency, Qext is computed using a Mie code
(Bohren, et al., 1998) with a log normal size
distribution, exploring the parameter space of
effective radius (Reff) from 0.1 to 2 micron, and
standard deviation (Veff) of 0.01 to 0.3 micron.
Given the Qext for each size distribution, the number
density ”n” is fitted using the relation β = n * Qext,
with β the extinction derived from the UVIS spectra.
The fit is made for each effective radius and each
standard deviation. The best fit finally selected will
be the one with the smallest reduced chi square.
The number density error is calcultated based on the
extinction error with a Monte Carlo algorithm.

Results:
Using only the spectral range of UVIS, the dust,
water ice and CO2 ice cannot be differentiated because the three aerosols have similar spectral features
in the UV-visible. However, it is possible to distinguish the particle size. Therefore, only dust will be
assumed in this work. Detection of CO2 and water ice
will be investigated in a future work
We observe three different spectral features characteristic of dust in the UV. Large particles above
1.2 micron show a relatively flat spectra, while oscillations seem to characterize the medium particles
between 0.5 and 1 micron, particles smaller than 0.5
micron are characterized by a spectral slope. These
features are illustrated in Figure 2.

superior at 1.2, we attribute a value of effective radius (Reff) of 1.5 µm and a standard deviation (Veff) of
0.2 as this is the most common mode observed
(Fedorova et al., 2014; Wolff, 2003).

Figure 3: Typical dust size vertical profile
Figure 3 shows an example of retrieved particle size
vertical profile. For this example, at low altitude and
up to 40km large particle are observed (Reff >1.2
micron). Above 40 km, a decrease in size, in agreement with the expected behavior of size decreasing
with altitude is observed.
Dust in the Martian atmosphere is sensitive to seasonal variations. During perihelion (LS 250), the
atmosphere of Mars becomes warmer, and dust can
be transported to higher altitudes. In the contrary, at
the aphelion (LS 70) dust remains confined at lower
altitudes.
This seasonal variability is shown in Figure 4

Figure 4: Dust vertical extinction profiles versus
solar longitude for Mars year 34 to 36

Figure 2 : Example of particles size effect on spectra
Because we are not sensitive to particles with size
larger than 1.2 micron. When the fit gives us a value

We can see on Figure 4 that at the perihelion dust is
present at higher altitudes and the extinction is
stronger than during the aphelion. In this work we
will further compare the vertical distribution of dust
for Mars year 34 (with global dust storm) and Mars
year 35 (without global dust strom), as well as
latitudinal variations.

Special case: detection of clouds

Figure 5: Detection of a cloud at high altitude
This characterization of size also helps to detect ice
clouds in the atmosphere. A cloud has been detected
for example in Figure 5 where we observe an increase of the particle size at high altitude around 70
km. Below that, the particles are smaller, the particles larger than 0.5 micron passed our selection criteria and are considered as real detection. To further
confirm the presence of the cloud, a comparison is
made with temperature profile derived from the
NOMAD SO channel (Trompet, L. et al., 2022). This
is illustrated in Figure 6.

Figure 6 : Comparisons between Temperature
profile and CO2 condensation temperature
We can see that for this particular case the retrieved
temperature goes below the CO2 condensation temperature around 70 km which supports the conclusion of the presence of a CO2 ice cloud.

Acknowledgements:
The NOMAD experiment is led by the Royal Belgian
Institute for Space Aeronomy (IASB-BIRA), assisted
by Co-PI teams from Spain (IAA-CSIC), Italy
(INAF-IAPS), and the United Kingdom (Open University). This project acknowledges funding by the
Belgian Science Policy Office (BELSPO), with the
financial and contractual coordination by the ESA
Prodex Office (PEA 4000103401, 4000121493), by
Spanish Ministry of Science and Innovation (MCIU)
and by European funds under grants PGC2018-

101836-B-I00
and
ESP2017-87143-R
(MINECO/FEDER), as well as by UK Space Agency
through grants ST/V002295/1, ST/V005332/1 and
ST/S00145X/1 and Italian Space Agency through
grant 2018-2-HH.0. This work was supported by the
Belgian Fonds de la Recherche Scientifique – FNRS
under grant number 30442502 (ET_HOME). The
IAA/CSIC team acknowledges financial support
from the State Agency for Research of the Spanish
MCIU through the ‘Center of Excellence Severo
Ochoa’ award for the Instituto de Astrofísica de Andalucía (SEV-2017-0709). US investigators were
supported by the National Aeronautics and Space
Administration. Canadian investigators were supported by the Canadian Space Agency. This project
has received funding from the European Union’s
Horizon 2020 research and innovation programme
under grant agreement No 101004052.

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