Water vapor (H2O) is only a trace species on Mars, with abundances of at most a few parts per thousand in an
otherwise carbon dioxide (CO2) dominated atmosphere. However, it plays a key role in Mars' atmospheric chem-
istry. Indeed, CO2 is subject to photolysis and its current abundances cannot be explained without assuming that
its photolysis product carbon monoxide (CO) is recycled into CO2 through a reaction with OH, a radical resulting
from photolysis of water vapor (McElroy & Donahue,1972; Parkinson & Hunten,1972). Another relatively
abundant trace gas on Mars is ozone (O3), but it is over a thousand times less abundant than water vapor on Mars,
and over a hundred times less abundant than O3 on Earth. Ozone is mainly produced by the three-body reaction
of atomic (O) and molecular (O2) oxygen in the presence of CO2. Yet fast reactions involving odd hydrogen
species (H, OH and HO2, or HOx) resulting from water photochemistry will readily destroy O3 or suppress its
formation in catalytic cycles, making water vapor the key species in controlling the abundance of O3. In the
lower atmosphere, there is sufficient CO2 (i.e., pressure) to maintain large O3 concentrations, but its abundance
Abstract The Nadir and Occultation for MArs Discovery (NOMAD)/UV-visible (UVIS) spectrometer on
the ExoMars Trace Gas Orbiter provided observations of ozone (O3) and water vapor in the global dust storm
of 2018. Here we show in detail, using advanced data filtering and chemical modeling, how Martian O3 in the
middle atmosphere was destroyed during the dust storm. In data taken exactly 1year later when no dust storm
occurred, the normal situation had been reestablished. The model simulates how water vapor is transported to
high altitudes and latitudes in the storm, where it photolyzes to form odd hydrogen species that catalyze O3.
O3 destruction is simulated at all latitudes and up to 100km, except near the surface where it increases. The
simulations also predict a strong increase in the photochemical production of atomic hydrogen in the middle
atmosphere, consistent with the enhanced hydrogen escape observed in the upper atmosphere during global dust
Plain Language Summary Global dust storms are rare but impactful events on Mars, occurring
about once in a decade. Previous investigations found how water vapor is redistributed throughout the entire
atmosphere in a dust storm. Photolysis of water vapor by sunlight produces highly reactive species that destroy
ozone (O3). Here we present O3 measurements taken by the NOMAD/UVIS instrument on the ExoMars Trace
Gas Orbiter in the 2018 global dust storm. After advanced data filtering, they demonstrate how O3 in the middle
atmosphere was much reduced compared to one Mars year later when no dust storm occurred. 3D atmospheric
model simulations of atmospheric chemistry in the global dust storm confirm this planet-wide O3 destruction,
and help to understand the involved processes. The simulations also predict a strong increase in production
of atomic hydrogen in the middle atmosphere, that can explain the observed increased hydrogen atmospheric
escape during global dust storms.
DAERDEN ET AL.
© 2022 The Authors.
This is an open access article under
the terms of the Creative Commons
which permits use, distribution and
reproduction in any medium, provided the
original work is properly cited and is not
used for commercial purposes.
Planet-Wide Ozone Destruction in the Middle Atmosphere on
Mars During Global Dust Storm
F. Daerden1 , L. Neary1 , M. J. Wolff2 , R. T. Clancy2 , F. Lefèvre3 ,
J. A. Whiteway4 , S. Viscardy1 , A. Piccialli1 , Y. Willame1, C. Depiesse1 , S. Aoki5 ,
I. R. Thomas1 , B. Ristic1 , J. Erwin1 , J.-C. Gérard6, B. J. Sandor2 , A. Khayat7 ,
M. D. Smith7, J. P. Mason8, M. R. Patel8 , G. L. Villanueva7 , G. Liuzzi7,9 ,
G. Bellucci10, J.-J. Lopez-Moreno11, and A. C. Vandaele1
1Royal Belgian Institute for Space Aeronomy BIRA-IASB, Brussels, Belgium, 2Space Science Institute, Boulder, CO, USA,
3LATMOS, Sorbonne Université, UVSQ Université Paris-Saclay, CNRS, Paris, France, 4Centre for Research in Earth and
Space Science, York University, Toronto, ON, Canada, 5Institute of Space and Astronautical Science ISAS, Japan Aerospace
Exploration Agency JAXA, Sagamihara, Japan, 6LPAP, STAR Institute, Université de Liège, Liège, Belgium, 7NASA
Goddard Space Flight Center, Greenbelt, MD, USA, 8School of Physical Sciences, The Open University, Milton Keynes, UK,
9Department of Physics, School of Arts and Sciences, American University, Washington, DC, USA, 10Istituto di Astrofisica e
Planetologia Spaziali, IAPS-INAF, Rome, Italy, 11Instituto de Astrofisica de Andalucia, IAA-CSIC, Granada, Spain
• NOMAD ozone (O3) data filtering
during the 2018 global dust storm
shows strong O3 destruction compared
to one year later with no dust storm
• 3D simulations of atmospheric
chemistry in the 2018 global dust
storm are presented to understand
impact on odd hydrogen and odd
• The model confirms middle-
atmospheric O3 destruction in the
dust storm and predicts increased
photochemical production of hydrogen
Supporting Information may be found in
the online version of this article.
Daerden, F., Neary, L., Wolff, M. J.,
Clancy, R. T., Lefèvre, F., Whiteway,
J. A., etal. (2022). Planet-wide ozone
destruction in the middle atmosphere
on Mars during global dust storm.
Geophysical Research Letters,
49, e2022GL098821. https://doi.
Received 22 MAR 2022
Accepted 20 MAY 2022
Conceptualization: F. Daerden
Data curation: F. Daerden, L. Neary, M.
J. Wolff, I. R. Thomas, B. Ristic
Formal analysis: F. Daerden, M. J. Wolff
Investigation: F. Daerden, M. J. Wolff,
R. T. Clancy, F. Lefèvre, J. A. Whiteway
Methodology: F. Daerden, M. J. Wolff
ExoMars Trace Gas Orbiter -
One Martian Year of Science
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decreases with height. This effect is strengthened by an increase in the photolysis rate coefficient of water vapor
with height. At the same time, water vapor frequently deposits into ice clouds below ∼40km (height changing
with latitude and season) and becomes only sparsely available above the cloud zone. In that region, the O3 forma-
tion is no longer suppressed, and its abundances will increase with height, although limited by the decreasing CO2
density (Daerden etal.,2019; Khayat etal.,2021; Lefèvre etal.,2004; Patel etal.,2021; Clancy and Nair,1996).
Until recently, O3 was observed mostly for its total column abundances (Lefèvre et al., 2021; Montmessin
etal.,2017; Perrier etal.,2006; Clancy etal.,2016; Willame etal.,2017) and in some cases for its vertical profile
(Lebonnois etal.,2006; Montmessin & Lefèvre,2013). The start of science operations of the ExoMars Trace Gas
Orbiter (TGO) in 2018 formed a landmark change in the knowledge of O3 on Mars. TGO carries two sensitive
spectrometer suites, Nadir and Occultation for MArs Discovery (NOMAD, Vandaele etal., 2018, 2019) and
Atmospheric Chemistry Suite (ACS, Korablev etal.,2018,2019). The UV-visible (UVIS) channel of NOMAD
monitors the vertical profile of O3 on a daily basis with the solar occultation technique (Khayat etal., 2021;
Patel et al., 2021). Water vapor profiles are measured simulateanously with the NOMAD IR channel (Aoki
etal.,2019). In addition, the IR channel of ACS could retrieve O3 in some cases (Olsen etal.,2020).
Dramatic changes in the vertical distribution of water vapor during global dust storms (GDS) were first reported
by the SPICAM instrument on Mars Express (Fedorova etal.,2018) for the 2007 GDS, and more recently by the
TGO (Aoki etal.,2019; Fedorova etal.,2020; Vandaele etal.,2019) and SPICAM (Fedorova etal.,2021) for
the 2018 GDS. As the planet's atmosphere contains much more dust, solar radiative heating by the dust changes
the temperatures considerably (Smith,2019) compared to the same season without a GDS. This has a dramatic
impact on water vapor, as the increased temperatures prevent cloud formation, and water vapor is transported
to much larger heights and, following the global circulation (which is also enhanced in dust storms), to higher
latitudes, up to the polar regions. This phenomenon was simulated in a General Circulation Model (GCM),
producing a good agreement with the water observations of NOMAD during the 2018 GDS (Neary etal.,2020).
The strong photochemical relationship between water vapor and O3 implies that the redistribution of water vapor
during the GDS will also have an impact on O3. A first analysis of NOMAD/UVIS O3 profiles taken during
the GDS (Patel etal.,2021) found that mid- and high-latitude O3 abundances above 20km were reduced in the
GDS compared to 1year later, when no GDS occurred. Here we perform new retrievals of the O3 profiles that
allow us to increase the signal-to-noise through spectral and spatial binning, and to develop an improved data
filtering. Both of these facets are particularly helpful for the lower transmittances found in dusty conditions, and
so to provide a clearer picture in the 2018 GDS. We will also investigate the behavior of odd hydrogen and odd
oxygen (including O3) using a GCM that is operated for the conditions of Martian years (MY) 34 and 35 (Daerden
etal.,2022; Neary etal.,2020) and that includes detailed atmospheric chemistry routines (Daerden etal.,2019).
Combining the filtered observations with the model simulations then allows to obtain a detailed picture of the
changes in atmospheric chemistry during the 2018 GDS.
2. Ozone Retrieval and Data Filtering
The NOMAD/UVIS spectrometer operates in the UVIS domain between wavelengths 200–650nm. It has two
observation modes with different sensitivity: the solar occultation mode and nadir/limb viewing mode. In this
study only observations made with the solar occultation channel are considered. Details and characteristics of the
NOMAD/UVIS instrument, its solar occultation channel, and calibration were presented in various publications
(Gérard etal.,2020; Khayat etal.,2021; Patel etal.,2017,2021; Vandaele etal.,2018).
We perform the retrievals of O3 profiles in a manner that is equivalent to previous work (Khayat etal.,2021; Patel
etal.,2021), but different in detail. First, we derive extinction profiles by applying the so-called “onion-peeling”
method to the NOMAD/UVIS transmittance spectra, as was described for TGO/ACS observations (Stcherbinine
etal.,2020). In essence, an upper triangular matrix is formed (at each wavelength) using a spherically symmetric
atmospheric shell model, which is solved analytically to produce vertical profiles of opacity with error propaga-
tion given by equation 7 of Stcherbinine etal.(2020). The column density of O3 in each layer is then determined
by fitting a model of a linear continuum plus O3 absorption (Sander etal., 2011) to the data between 240 and
320nm. The best-fit solution is found using a Marquardt-Levenberg algorithm, as implemented in the package
MPFIT (Markwardt,2009). The uncertainty associated with the derived O3 densities (and linear fit parame-
ters) is provided by MPFIT from the diagonal terms of the covariance matrix. Extensive comparisons of our
Resources: F. Daerden, L. Neary, M. J.
Wolff, I. R. Thomas, B. Ristic
Software: F. Daerden, L. Neary, M. J.
Supervision: F. Daerden
Validation: F. Daerden, M. J. Wolff
Visualization: F. Daerden, L. Neary, R.
T. Clancy, F. Lefèvre
Writing – original draft: F. Daerden
Writing – review & editing: F. Daerden,
L. Neary, M. J. Wolff, R. T. Clancy, F.
Lefèvre, J. A. Whiteway, J. Erwin, J.-C.
Gérard, B. J. Sandor, M. R. Patel
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retrievals outside of dusty conditions with the published full NOMAD/UVIS O3 data set (Khayat etal.,2021;
Patel etal.,2021) showed good agreement within the uncertainties of each retrieval. The data set presented in this
paper, with full retrieval and filtering (see below) details, is publicly available (Daerden & Wolff,2022).
The motivation to develop our own retrieval process was the ability to perform sensitivity analyses and develop
additional filters to remove spurious O3 detection. Such spurious detections (i.e., O3 detection without an O3
signature present in the transmission spectra) were found to occur in regimes where the opacity spectra can be
quite noisy. In other words, relying on the covariance matrix alone for our retrievals was not sufficient, since it can
produce unphysical results. In this direction, we have taken two approaches. In addition to the formal uncertainty
returned by MPFIT, we calculate an empirical equivalent to the information content approach (Rodgers,2000) by
comparing the χ
2 of the O3 retrieval for each altitude to that of a linear-only model such that values less than one
indicate that allowing O3 provides a better fit, that is, m=χ
2(linear)<1. In other words, if a linear
model fits as well (or better) than the O3 model, the retrieved value is not statistically significant. We also allow
for the binning of transmittance data in both wavelength and height, where the latter dimension is particularly
effective in reducing the noise in the O3 retrievals when the observations employ a large number of positions per
occultation. We generally employed bins of 10nm and 4km. A somewhat similar analysis to ours presenting an
information content approach to the full NOMAD/UVIS O3 data set is in preparation by A. Piccialli etal.
Through numerical experimentation, we developed the following data filtering criteria. Data are removed from a
profile if (a) the average transmittance between 240 and 320nm was below 0.02, and (b) when m=χ
2(linear)>0.7, except during the period of April–November 2018 (which suffered from reduced signal-to-
noise ratio) where we use m>0.85. Outside of this period, the value of 0.7 is preferred since it provides a more
consistent removal of spurious data points. Both m values are the result of substantial trial-and-error processes
which include manual inspection of many (e.g., hundreds) individual profile fits per experiment. (c) A final step
in the filtering process is done when the retrieved O3 number density was less than the uncertainty returned by
3. Ozone in the Global Dust Storm
The observations are shown in Figure1 for latitude bands >45°N and <45°S respectively, for the period before
and during the 2018 GDS (in Mars year, MY, 34) and in the same season exactly one Martian year later (MY35),
when no GDS occurred. The latitude of the observations changes over time, and their geospatial distribution in
both years is shown in Figure S1 inSupporting InformationS1. The choice to show the latitude bands poleward
of 45° resulted from a compromise to minimize the mixing of data taken at different latitudes while maintain-
ing a sufficiently dense temporal coverage of the profiles. Figure2 shows the observations for all latitudes, but
averaged over wider Ls intervals. The data are plotted using the height above the MOLA Mars reference level
(Smith etal.,1999) as vertical coordinate, in order to combine observations taken over a variety of latitudes and
longitudes (and hence for different surface heights). The observation errors for the data shown in both figures are
shown in Figures S2 and S3 in Supporting InformationS1.
In the period before the 2018 GDS (Ls=160°–187°, Figure1 and Figure2a,b), a distinct O3 minimum between
30 and 50km (height varying with time and latitude) can be seen in each hemisphere at high latitudes in both
years, with O3 remaining abundant above this minimum up to 50–60km (height again varying with time and
latitude, this is the high altitude O3 layer that was discussed before, Khayat etal.,2021). This pattern is present in
both Mars years, but in MY34, at southern latitudes, it is less apparent because of the sparse observations in this
case (Figure S1 in Supporting InformationS1). The minimum in O3 is associated with an observed and simulated
maximum in water vapor in this region that is caused by transport of wet air along the ascending branches of the
Hadley circulation cells (Aoki etal.,2019; Daerden etal.,2019; Neary etal.,2020). At these heights, water is
more easily photolyzed and produces odd hydrogen species that both destroy and suppress O3 formation (Daerden
Observational constraints imposed by the orbital geometry of the TGO spacecraft caused gaps in the data of up to
∼15° wide in Ls (at different times of the year in MY34 and MY35, see Figure1), complicating the comparison of
the 2years over certain Ls ranges. Nevertheless, it is seen that in MY35 the ozone peak abundances above 30km
continue through to Ls=230° after which the observable ozone above 30km gradually disappears (Figures1b
and1d). Figure2d confirms this seasonal behavior in MY35, with a similar ozone distribution as before Ls=187°
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Figure 1. Time series of the ozone (O3) number density profiles. Panels (a and b) for latitudes north of 45°N, for the same season in Mars year 34 (a, with GDS) and in
Mars year 35 (b). Panels (c and d) show the same for latitudes south of 45°S. Left plots for the O3 observed by NOMAD/UVIS and right ones from simulations. Gaps in
the data are due to observational restrictions imposed by orbital geometry of the Trace Gas Orbiter spacecraft. Gray shading: for the observations, this represents values
that were filtered out (see Section2) or are below 10
−3; for the simulations, values are below 10
−3. Height is taken with respect to the MOLA reference level
(black shading at the bottom of the panels shows the actual surface). Model results were interpolated to the time and location of the observations. The thick vertical
dashed lines indicate the onset of the global dust storm in MY34 and the same time one Mars year later.
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Figure 2. Latitude-height distribution of the ozone (O3) number density profiles. Panels (a and b) show the average over Ls=160°–187° and (c and d) over Ls=187°–
230°, in MY34 (a and c) and MY35 (b and d). The first period is before the global dust storm (GDS) in MY34, the second period covers the GDS in MY34. Left plots
for the O3 observed by NOMAD/UVIS and right ones from simulations. Gray shading: for the observations, this represents values that were filtered out (see Section2);
for the simulations, values are below 2×10
−3. Height is with respect to the MOLA reference level (black shading shows zonally averaged MOLA topography). A
grid of 1.5° in latitude and 1.5km in height was used to compute the averages. Model results were interpolated to the time and location of the observations.
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(Figure2b) but with slightly reduced abundances. In MY34 however, ozone is strongly reduced immediately after
the onset of the GDS for all heights above 20km in the north and 30km in the south (Figures1a and1c), consist-
ent with the previous analyses (Khayat etal.,2021; Patel etal.,2021). This dramatic ozone reduction takes place
at all latitudes, as can be seen in Figure2c.
4. Atmospheric Chemistry Simulations
The differences in O3 behavior between MY34 and MY35 have been qualitatively attributed to the observed
increase of water vapor during the GDS at heights and latitudes that, in normal (non-GDS) conditions exhibit
very low water vapor abundances (Patel etal.,2021). The enhanced high-altitude water vapor abundances lead
to increased water vapor photolysis, producing odd hydrogen species that rapidly catalyze O3 loss. However, the
enhanced odd hydrogen species that are key to this process were not observable at the time. Without observational
access to these radical species that drive the chemical pathways between water vapor and O3, we turn to a model
for atmospheric chemistry (Daerden etal.,2019; Neary & Daerden,2018) to understand the enhanced destruction
of O3 during the GDS. Previous work has demonstrated how this model reproduced the redistribution of water
vapor observed by NOMAD during the 2018 GDS (Neary etal.,2020). Here the model was run consecutively
for MY34 and MY35. Both years were simulated using their respective daily dust optical depth climatologies
(Montabone etal., 2015, 2020). For a direct comparison of model and data, the time-evolved 3D simulation
results were interpolated in space and time to the NOMAD/UVIS observed O3 profiles and added to Figures1
and2 in the same format as the data, but without filtering.
The resemblance of the model results to the O3 observations is strong in morphology and often also in absolute
values, providing evidence that the model captures well the underlying chemical processes. To understand these
processes, Figure3 shows the time-evolved simulation results of a range of key species (H2O, OH, H, HO2, O and
O3) at latitudes 60°N and 60°S in both Mars years. Figure4 shows the global latitude-height cross-sections of
the same species averaged over Ls=210°–220° (peak of the GDS in MY34) in both MY34 and MY35, and their
relative change in this time window. (As the changes for odd hydrogen species between the 2years are very large
(ratio >1,000), we show them as a ratio. For O3, the differences are smaller (ratio ∼< 20), and we show them as a
percentage difference on a linear scale.) Both figures assist to the interpretation of Figures1 and 2.
In the middle atmosphere, up to 100km, and as low as 20km at at low- and midlatitudes, down to 10km north
of of 45°N latitude, and down to the surface south 60°S, the O3 abundances are severely reduced, up to 100%
(Figure 4f, right panel, and Figure S8 in Supporting Information S1, right panels). This is a planet-wide O3
destruction, taking place throughout almost the entire atmosphere, and demonstrates the massive impact of the
GDS on the atmospheric composition.
Because water vapor is redistributed in a GDS from the lower atmosphere at low latitudes to the rest of the
atmosphere (Neary etal.,2020), a small decrease in water vapor columns could be seen in measurements during
previous GDS (M. D. Smith,2004; M. Smith etal.,2018; Trokhimovskiy etal.,2015). As a consequence, we here
simulate that O3 abundances at low altitudes, low- and midlatitudes are increased (Figure4f).
In terms of total O3 columns, these changes are small in absolute terms, because (a) the air densities in the lowest
atmosphere dominate the total column contribution, and (b) because total O3 abundances in this season are
small. But the changes are not small in relative terms. Figure S1 in Supporting InformationS1 shows how the
O3 column at low latitudes is predicted to increase from below 1 to almost 2 μm-atm, this is a 100% increase. At
high latitudes, the reduction of the O3 column is up to 2 μm-atm, representing a 50% decrease in the north and a
100% decrease in the south.
Five reactions (see below) are dominating the change of the O3 in the GDS. This change in O3 follows the change
in the relative strength of these reactions during the GDS (Figures S4 through S7 in Supporting Information S1
for reaction numbering and rates). Combined, they define the main pathway from enhanced water vapor to O3
depletion, which can be described as follows:
1. The increase of water vapor in the middle and upper atmosphere and its subsequent photolysis (H2O+hν →
H+OH, J11-13) results in the increased formation of atomic hydrogen (H) and OH radicals.
2. The OH radicals react with CO (CO+OH → CO2+H, R1), and to a lesser extent with O (O+OH → O2+H,
R10), to cause an additional increase in atomic hydrogen.
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3. Atomic hydrogen reacts with molecular oxygen (H+O2+CO2 → HO2+CO2, R26) leading to an increase
4. HO2 then reacts with atomic oxygen, O (HO2+O → OH+O2, R11) and so reduces the abundance of atomic
oxygen. OH is returned to the atmosphere to resume step 2, and acts as a catalyzer.
5. Finally, the decrease in O suppresses the formation of O3 in O+O2+CO2 → O3+CO2 (R24)
The direct destruction of O3 by HO2 and OH is also enhanced in the GDS (Figure S7 in Supporting Informa-
tionS1, R18 and R21), but only at high altitudes/latitudes, and not in the region of highest O3 abundances, and so
contributes less to the O3 changes in the GDS.
The simulated time evolution of O3 throughout the GDS as a result of these reactions, in comparison with the
simulated evolution in MY35 and with the relative differences between both years, is shown Figure S8 in Support-
ing InformationS1. An animation showing the model simulation of the O3 vertical distribution in the GDS and
one year later is also included in the Supporting InformationS1.
A question to consider is to what extent the specific spatio-temporal distribution of the NOMAD observations
(Figure S1 in Supporting InformationS1) could have an impact on our results. In some cases, there are differ-
ences in coverage between MY34 and MY35, such as in the south before the GDS, or in the northern high-lati-
tudes after Ls∼200°. However between Ls∼215°–230°, the coverage is similar in both years, so that the results
shown in Figure2 are robust. When considering only the profiles that are close in both Ls and latitude in both
Figure 3. Simulated time series of H2O, OH, H, HO2, H and O3 number densities for latitudes 60°N (top rows) and 60°S (bottom rows) around equinox, for Martian
years 34 (with global dust storm,GDS) and 35 (no GDS). The color shading shows the number density averaged over all longitudes and local times. The dashed white
lines indicate the onset and end of the GDS in MY34 and the same times in MY35.
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Figure 4. Latitude-height cross-sections of the simulated number densities of H2O, OH, H, HO2, O and O3 for MY34 (left
column) and MY35 (center column), averaged over Ls=210°–220° (period which falls in the global dust storm (GDS) in
MY34). The right column show the ratio of the averaged number densities in MY34 and MY35, except for O3 for which
the relative difference is shown. The densities from both years were first interpolated to a common altitude grid as the
atmospheric height scale changed during the GDS.
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years (Figures S9 and S10 in Supporting InformationS1), there are more gaps in the figures, but they confirm our
conclusions about how NOMAD witnessed O3 loss in the 2018 GDS.
An interesting result from the simulations is the increased production of atomic hydrogen in the middle and
upper atmosphere (Figures3c and4c), with increased number densities by factors of 10–100 above 60km. This
result (already preliminary shown in Neary etal.,2020) seems consistent with the observed enhanced abundance
of atomic hydrogen in the upper atmosphere as observed by MAVEN during the 2018 global and regional dust
storms (Chaffin etal.,2021; Stone etal.,2020).
The two endpoints of the above chain of reactions, water vapor and O3, were observed by NOMAD before and
during (MY34) and out of (MY35) the 2018 GDS (Aoki etal.,2019; Khayat etal.,2021; Patel etal.,2021). The
model demonstrates the detailed chain of reactions occurring between these end points, involving unobserved
odd hydrogen and oxygen species. The successful reproduction of the water vapor (Neary etal.,2020) and O3
observations (this paper)—both end points in the chemistry chain—in very different conditions (in and out of a
GDS), by the model provides support for the simulation of these unobserved species and the involved reactions.
The simulated O3 destruction in the GDS, supported by available observations, occurs throughout the middle
atmosphere, that is, on a planet-wide scale. This shows the massive impact of the GDS and the redistribution of
water vapor on atmospheric chemistry and composition, as the odd hydrogen radicals resulting from water vapor
photolysis are highly reactive.
Interestingly, the simulated increase in the production of atomic hydrogen in the middle atmosphere during the
GDS (Figures 3c and 4c) is important to understand the observed enhanced atmospheric escape during dust
storms (Chaffin etal.,2021; Stone etal.,2020). It has been generally assumed that the redistribution of water
vapor during GDS is causing the observed enhanced escape (Chaffin etal.,2021; Heavens etal.,2018; Shaposh-
nikov etal.,2021; Stone etal.,2020). While a dynamical perspective to understand the transport of water vapor
to the upper atmosphere during a GDS was given in terms of wave activity (Yiğit,2021), we provide here the
photochemical processes that are also involved.
Data Availability Statement
ExoMars Trace Gas Orbiter data are publicly available through the European Space Agency's Planetary Science
Archive (http://archives.esac.esa.int/psa) with additional access to NOMAD data through the PI institute (http://
nomad.aeronomie.be). The NOMAD/UVIS Solar Occultation O3 data set presented in this work, as well as the
results from the General Circulation Model simulations are available on the BIRA-IASB data repository (https://
repository.aeronomie.be/?doi=10.18758/71021070, Daerden & Wolff, 2022). The GEM-Mars General Circu-
lation Model is based on the Global Environmental Multiscale model 4.2.0 version of the community weather
forecasting model for Earth, which is one of the more recent versions available to the community, under the GNU
Lesser General Public Licence v2.1. The adaptation for Mars is developed and maintained at the Royal Belgian
Institute for Space Aeronomy.
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storms observed by TGO/NOMAD. Journal of Geophysical Research: Planets, 124(12), 3482–3497. https://doi.org/10.1029/2019JE006109
Chaffin, M. S., Kass, D. M., Aoki, S., Fedorova, A. A., Deighan, J., Connour, K., etal. (2021). Martian water loss to space enhanced by regional
dust storms. Nature Astronomy, 5(10), 1036–1042. https://doi.org/10.1038/s41550-021-01425-w
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This work was made possible thanks
to the reconstructed gridded maps of
column dust optical depth from Mars
Climate Sounder observations provided
by L. Montabone. The dust maps were
prepared using MCS v5.3 provided by A.
Kleinböhl and D. Kass. Dust clima-
tologies can be found at the following
mars/dust_climatology/. ExoMars is a
space mission of the European Space
Agency (ESA) and Roscosmos. 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 the UK Space Agency
(grants ST/V002295/1, ST/P001262/1,
ST/V005332/1 and ST/S00145X/1), by
the Spanish Ministry of Science and
Innovation (MCIU), and by European
funds (grants PGC2018-101836-B-I00
and ESP2017-87143-R, MINECO/
FEDER), as well as by the Italian Space
Agency (Grant 2018-2-HH.0). This work
was supported by the Belgian Fonds de
la Recherche Scientifique – FNRS (Grant
Nos. 30442502, ET_HOME). This work
has received funding from the European
Union's Horizon 2020 research and
innovation programme (grant agreement
No 101004052, RoadMap project). 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 inves-
tigators were supported by the National
Aeronautics and Space Administration,
by NASA's Mars Program Office (under
WBS 604796, “Participation in the TGO/
NOMAD Investigation of Trace Gases on
Mars.”), and by NASA (award number
80GSFC21M0002). Canadian investi-
gators were supported by the Canadian
Space Agency. We thank Manuel López
Puertas and two anonymous reviewers for
useful comments on the manuscript.
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