Water vapor mapping on mars using omega mars express

Planetary and Space Science 55 (2007) 333–342
Water vapor mapping on Mars using OMEGA/Mars Express
R. Melchiorria
, T. Encrenaza,Ã, T. Foucheta
, P. Drossarta
, E. Lelloucha
, B. Gondetb
,
J.-P. Bibringb
, Y. Langevinb
, B. Schmittc
, D. Titovd
, N. Ignatieve
a
LESIA, Observatoire de Paris, France
b
IAS, Orsay, France
c
LPG, Grenoble, France
d
MPI, Lindau, Germany
e
IKI, Moscow, Russia
Accepted 30 May 2006
Available online 1 September 2006
Abstract
A systematic mapping of water vapor on Mars has been achieved using the imaging spectrometer OMEGA aboard the Mars Express
spacecraft, using the depth of the 2.6 mm (n1, n3) band of H2O. We report results obtained during two periods: (1) Ls ¼ 330–401
(January–June 2004), before and after the equinox, and (2) Ls ¼ 90–1251, which correspond to early northern summer. At low latitude,
our results are globally consistent with previous measurements from ground-based and space (MAWD/Viking and TES/MGS)
observations. However, at early northern summer and at high northern latitude (70–80 1N), the water vapor abundances, which we
retrieved, appear to be weaker than MAWD and TES results. At the time of water sublimation during early northern summer, there is a
maximum of water vapor content at latitudes 75–801N and longitudes 210–241E. This region is not far from the area where OMEGA
identified a high abundance of calcium-rich sulfates, most likely gypsum. Our data provide the first high-resolution map of the martian
water vapor content above the northern polar cap.
r 2006 Elsevier Ltd. All rights reserved.
Keywords: Mars; Mars atmosphere; Infrared spectroscopy
1. Introduction
The water vapor cycle on Mars is known to be very
complex, and is a precious indicator of the planet’s climatic
variations. It has been regularly monitored over several
decades, using ground-based observations and space
experiments, in order to better understand the seasonal
and interannual variations of water vapor on Mars.
Basically, two types of spectroscopic data have been used:
the near-infrared H2O bands in the reflected part of the
spectrum, and the rotational water lines, beyond 20 mm, in
the thermal infrared. The first global thermal mapping of
H2O was performed by the IRIS interferometer aboard
Mariner 9 (Conrath et al., 1973). It was followed by the
very complete survey achieved at 1.38 mm by the MAWD
(Mars Atmosphere Water Detection) experiment aboard
the Viking orbiters; these results were used for decades as a
reference (Farmer et al., 1977; Jakosky and Farmer, 1982;
Fedorova et al., 2004). In parallel, ground-based monitor-
ing of Martian water vapor was achieved, taking advantage
of the Earth–Mars water line Doppler shift or using
isotopic HDO lines. Observations were performed in the
near infrared range (Sprague et al., 1996, 2003), in the
millimeter range (Encrenaz et al., 1991, 1995, 2001) and in
the radio range (Clancy et al., 1992, 1996). In addition, the
water vapor abundance was also derived from the H2O
bands at 2.6 mm (n1, n3) and 6.2 mm (n2) using the Short-
Wavelength Spectrograph (SWS) instrument aboard the
Infrared Space Observatory (ISO) Earth-orbiting satellite
(Lellouch et al., 2000). Local measurements were obtained
at the same time (June–August 1997) by the Mars
Pathfinder lander (Titov et al., 1999). More recently, the
TES instrument aboard the Mars Global Surveyor mission
has achieved a complete coverage of the water vapor
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E-mail address: therese.encrenaz@obspm.fr (T. Encrenaz).

column density, using the rotational H2O transitions at
20–35 mm (Smith, 2002, 2004).
All these data sets have shown evidence for large spatial
and temporal variations of the Martian water vapor
abundance, with a pronounced maximum, as high as
80 pr-mm, at high northern latitudes at the time of early
northern summer (Ls ¼ 90–1201), while minima in the
southern hemisphere at the same time are as low as 1–3
pr-mm. There is a noticeable asymmetry between the
northern and southern hemispheres, as the maximum at
high southern latitudes, at the time of early southern
summer, is only about 30 pr-mm. This asymmetry has been
explained by the eccentricity of Mars which is such that
northern summer takes place near aphelion, leading to a
temperature difference of about 20 K between maximum
temperatures at high latitudes during northern and south-
ern summers (Clancy et al., 1996). Global climatic models
(Forget et al., 1999; Montmessin et al., 2004) are able to
account for the general behavior of the seasonal water cycle
as a result of seasonal cycle of Mars, due to the periodic
condensation of CO2 and H2O at the polar caps. However,
the amplitude of the asymmetry is not fully reproduced,
and the exact abundance of the water maximum at early
northern summer is not well constrained by climatic
models (Forget et al., 1999).
The OMEGA imaging spectrometer aboard the Mars
Express spacecraft offers a new opportunity to monitor the
water vapor abundance on Mars, through the analysis of
the (n1, n3) band at 2.6 mm. This band, in the wing of the
strong (n1+n3) CO2 band, centered at 2.75 mm, is located in
the reflected solar part of the Martian spectrum. Thus, it
has the advantage of being, to first order, independent of
the temperature profile. In contrast, the analysis of spectral
signatures at higher wavelengths (n2 at 6.2 mm, rotational
lines beyond 20 mm), formed in the thermal regime, requires
a good knowledge of the thermal profile. We note,
however, that the use of the reflected solar component
leads to some uncertainty due to scattering by martian
aerosols; this effect is discussed in more detail in Section 3.1.
Another advantage of the reflected sunlight analysis is that
it allows to retrieve the water vapor content everywhere on
the Martian disk, even over the polar caps where the
surface temperature is too low for the thermal flux to be
detectable.
In this paper, we present the water vapor mapping
obtained during the first year (2004) of the Mars Express
operation. Section 2 presents the instrument and the
observations. Section 3 describes the method used, and
the atmospheric modelling. Section 4 shows the results
obtained (1) before and after the north hemisphere spring
equinox (Ls ¼ 330–401) when the water content is expected
to be rather low and more or less uniform, and (2) at high
northern latitudes during early northern summer, when the
water vapor content is maximum, due to the sublimation of
the northern ice cap (the whole set of maps are East
longitude oriented). Preliminary results of this study have
been published in Encrenaz et al. (2005).
2. The observations
The Mars Express spacecraft (Chicarro et al., 2004) was
launched by ESA on June 2, 2003, and has been operating
in orbit around Mars since January 2004. The orbit is
almost polar and highly elliptical, with a period close to 7 h
(6.72 h for the 100 first days, and 7.58 h after). As a result, a
given location observed by Mars Express on a given orbit
was again observed, with a slight shift, 11 orbits later (or 13
orbits during the first 100 days). The periapsis was
processing in latitude by about 201 per month. The latitude
of the periapsis started from about 01 at the beginning of
the mission, decreased down to about 80S in June 2004
(orbit 500) then increased again regularly to reach a
maximum latitude of about 8 1N in March 2005 (orbit
1500), and so on.
The OMEGA instrument (Bibring et al., 2004) is an
imaging spectrometer operating in the visible and near-
infrared range, from 0.35 to 5.1 mm, with a spectral
resolution of 7 nm below 1 mm, 14 mm in the 1.0–2.5 mm
range, and 20 nm above 2.5 mm. With an instantaneous
field of view (IFOV) of 1.2 mrad, its spatial resolution at
the surface of Mars ranges from about 300 m (close to
periapsis) up to 4.8 km from an altitude of 4000 km. In
order to have a continuous coverage during an orbit, the
longitudinal width varies from 16 to 128 IFOV, depending
on the distance and on the speed of the instrument from the
surface. For this reason each orbit is divided in several
sessions; namely an orbit is defined by a code xxxx-y where
x is the number of the orbit since the beginning of the
mission and y is the session associated (sessions do not
define a pixel width). The present analysis is based on the
data of the IR-short-wavelength channel of OMEGA,
which extends from 1.0 to 2.7 mm.
The data presented here cover two periods:
Orbits 6–520 (January 2003–June 2003), Ls ¼ 330–401.
These data cover mostly the mid-latitude regions, at a
season where the water vapor content is expected to be
low (Jakosky and Farmer, 1982; Smith, 2002, 2004),
with little temporal evolution.
Orbits 900–1150 (October 2003–January 2004),
Ls ¼ 90–1251. These data cover the high northern
latitudes at the time of northern spring, when the water
vapor content is expected to be at the maximum due to
the sublimation of the water ice cap.
Fig. 1 shows two raw OMEGA spectra between 1.0 and
2.7 mm, recorded during Orbit 0037_3 , at the summit of
Olympus Mons (lat. ¼ 301N, long. 2271E) and at the foot
of the volcano (lat. ¼ 171N, long. 2271E). Each spectrum is
the average of 352 individual spectra, corresponding to 11
pixels along the latitude (spacecraft motion) and 32 pixels
along the longitude (for a surface of 0.81 lon. Â 0.11 lat.).
All OMEGA data were flat-fielded, corrected from non-
linearity, dark current and bad pixels, and calibrated using
a standard procedure.
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R. Melchiorri et al. / Planetary and Space Science 55 (2007) 333–342334

In Fig. 1, all differences between the spectra at the
summit and at the foot of Olympus Mons are due to the
different intensities of atmospheric features (Bibring et al.,
1991). Indeed, the mineralogy is basically the same at both
places, and the only difference is the atmospheric path
between the spacecraft and the surface: the altitude at the
summit is more than 20 km above the reference level while
our foot spectrum is about 1 km below it. The shape of
atmospheric features is better illustrated in Fig. 2, which
shows the ratio of the two previous spectra (Foot/Summit).
As the atmospheric path is by far longer for the foot
spectra than that at the summit, the atmospheric signatures
appear in absorption in the ratio.
In order to analyze these features, we have divided all
individual spectra by the averaged summit spectrum shown
in Fig. 1. This allows us to eliminate the uncertainty
associated with the OMEGA instrumental transfer func-
tion and to obtain a good fit of observed atmospheric
bands with modelled spectra. Because of the great stability
of the instrumental parameters since the beginning of the
operations, this method can be applied on the whole data
set (at least up to orbit 1500), although it covers a wide
temporal range (Mustard et al., 2005).
3. The radiative transfer model and the analysis of the
2.6 lm water band
3.1. Atmospheric and spectroscopic parameters
Comparing the ratio of observed spectra with synthetic
calculations first requires a proper modelling of the
reference spectrum corresponding to the summit of
Olympus Mons. Using the European Martian Climate
Database (EMCD) developed at Oxford and at the
Laboratoire de Me´ te´ orologie Dynamique (LMD) in Paris
(Forget et al., 1999), we have chosen the atmospheric
parameters corresponding to this location, for the seasonal
period (Ls ¼ 3371) and the local time (13:00) correspond-
ing to the OMEGA observation of Olympus Mons. We
used a surface pressure of 1.1 mbar and a surface
temperature of 275 K. The atmospheric temperature
decreased from 190 K at the surface down to 150 K at
P ¼ 0.01 mbar (z ¼ 45 km) and staid constant above this
altitude level. Note that there is a strong temperature
contrast (85 K) between the surface temperature at
Olympus summit and the atmospheric temperature at
z ¼ 0 just above it. At the foot of Olympus Mons, we
assumed a surface temperature of 280 K and a temperature
profile ranging from 250 K at P ¼ 7.5 mbar to 200 K at
1.6 mbar and 150 K at P ¼ 0.01 mbar. In this case, the
temperature contrast between the surface and the first
atmospheric layer is 25 K. We kept the surface pressure at
the Olympus foot as a free parameter to be determined by
the best fit of the CO2 bands.
Radiative transfer calculations over the 1.0–2.7 mm range
were performed using a line-by-line code including the
spectral signatures of CO2, H2O and CO. Spectroscopic
parameters were taken from the GEISA data bank
(Jacquinet-Husson et al., 1999). Line-broadening para-
meters for CO2–CO2, H2O–CO2 and CO–CO2 collisions
were taken from Pollack et al. (1993) and references
therein. We used a frequency step of 0.01 cmÀ1
, which was
found to be small enough for the spectral signatures to be
accurately reproduced. The results were convolved with the
OMEGA spectral instrumental function.
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0.30
0.25
0.20
0.15
0.10
I/F
0.05
0.00
1.0 1.5 2.0
wavelength (µm)
2.5
Fig. 1. Examples of two averaged I/F spectra, at the summit ((lat ¼ 171N,
long ¼ 227E, red curve) and at the foot (lat ¼ 301N, long ¼ 2271E, black
curve) of Olympus Mons (Orbit 0037_3). Each spectrum is the average of
352 individual spectra corresponding to 11 pixels along the latitude
(spacecraft motion) and 32 pixels along the longitude. The main spectral
features are atmospheric features, which are much weaker at the top of the
volcano (red spectrum) than at the bottom (black spectrum).’’
1.1
1
0.9
0.8
0.7
0.6
0.5
0.4
0.3
1 1.2
Spectrumratio(OlympusFoot/Summit)
1.4 1.6 1.8
Wavelength (micrometers)
2 2.2 2.4 2.6
Fig. 2. Ratio of the two spectra shown in Fig. 1. Color curves: synthetic
ratios calculated with averaged surface pressures of 7.5 mbar at the foot
and 1.1 mbar at the summit. A constant mixing ratio is assumed for
Q(H2O): 150 ppm (red), 300 ppm (green) and 600 ppm (blue). The best fit is
achieved for Q(H2O) ¼ 150 ppm.
R. Melchiorri et al. / Planetary and Space Science 55 (2007) 333–342 335

As mentioned in the introduction, our calculations do
not include scattering by aerosols; indeed, implementing
this procedure in our calculations would be, in terms of
computer time, incompatible with the automatic method
developed in the present study. Scattering by dust may
have two opposite effects on our results. On one hand,
multiple scattering can enhance the path length and thus
lead to an overestimate of the gaseous column density. On
the other hand, a layer of aerosols may reflect the solar
photons above the surface, and thus reduce the observed
column density. The latter effect is sometimes observed in
the core of the strong CO2 band at 2.0 mm. However, it can
be seen that, even in this case, the depth of the weak CO2
bands (at 1.44 mm in particular) is unaffected. We can thus
conclude that the effect of dust scattering is, to first order,
negligible in the case of the 2.6 mm H2O band, which has a
maximum depth of a few percent.
3.2. Modelling the reference spectrum
We have first determined the surface pressure corre-
sponding to our Olympus foot spectrum by fitting the CO2
bands, which shows a series of bands of different
intensities. The strongest one appears at 2.75 mm (n1+n3),
then the next one (2n1+n3) is visible around 2.0 mm, and
weaker ones are present at 1.6 mm (3n1+n2) and 1.44 mm
(3n3). It can be seen that a very good fit is obtained over the
whole spectral range when a surface pressure of 7.5 mbar is
used at the Olympus foot, which is consistent with the
GCM predictions, based upon an interpolation of the
MOLA data ($7.3 mbar).
Fig. 2 also shows 3 synthetic models corresponding to 3
different values of the H2O mixing ratio, assumed to be
constant with altitude over the flank of Olympus Mons
(Q(H2O) ¼ 150 ppm, 300 and 600 ppm). Water vapor
signatures are present at 1.38 mm (n1+n3) and 2.6 mm
(n1, n3), in the wing of the strong 2.7 mm CO2 band. We
have selected the 2.6 mm wavelength for our analysis (on
the slope of the 2.7 mm band), as it is by far the strongest
signature, and also because the 2.6 mm range is free of
mineralogical signatures. Fig. 2 shows that the best fit of
the water band is obtained with an H2O mixing ratio of
150 ppm. We have thus assumed this value for modelling
our reference spectrum of Olympus summit; the corre-
sponding H2O column density is 1.6 pr-mm. However,
it should be mentioned that there is some uncertainty
associated with the hypothesis of a constant H2O mixing
ratio along the flanks of Olympus Mons. We choose
this assumption as the simplest one. However, if the
water-mixing ratio at the top of Olympus was twice its
value at the foot, then the water column density would be
3.2 pr-mm, while it could be as low as zero if the water vapor
was significantly depleted at the summit. Assuming that the
H2O mixing ratio at Olympus summit is not enhanced by
more than a factor 2, we infer that the H2O column density
at the summit is 272 pr-mm. The corresponding error, on
all measurements, associated to this uncertainty, is 2 pr-mm.
3.3. Analysis of the 2.6 mm H2O band
Synthetic calculations show that the depth of the 2.6 mm
water band depends not only upon the H2O column
density, but also upon the surface pressure while it is, over
a wide range, independent upon the temperature profile.
Then, we have built a curve of growths of the 2.6 mm H2O
band depth as a function of the H2O surface mixing ratio
Q, for a given set of surface pressures Ps. These curves of
growth are shown in Fig. 3a for a set of 4 surface pressures
(1.4, 4.0, 7.5 and 10.0 mbar). We have checked that these
curves of growth are independent of the temperature over a
large range of atmospheric profiles, as the measured H2O
band depth varies by a few percent at most.
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0.12
0.1
0.08
0.06
0.04
0.02
H2ObanddepthH2Ocolumndensity(pr-microns)
0
0 20 40 60 80 100 120 140
water partical pressure (Q x ps x 10-7 bar)
160
140
120
100
80
60
40
20
0
0 0.002 0.004 0.006
H2O surface mixing ratio
0.008 0.01
(a)
(b)
Fig. 3. (a) Curves of growth of the 2.6 mm H2O band, calculated for
different values of the surface pressure. From right to left : Ps ¼ 1.4, 4.0,
7.5 and 10.0 mbar. The abscissa is the product of the surface pressure by
the H2O mixing ratio, i.e. the H2O partial pressure at the surface. The
ordinate is the depth of the water band: D ¼ 1À[2 Â I(2.60 mm)/
(I(2.55 mm)+I(2.65 mm))]. In absence of saturation, this quantity is
proportional to the H2O column density. (b) The water vapor column
density as a function of the H2O mixing ratio, for different values of the
surface pressure, calculated for a cold temperature profile, for which the
saturation effect is maximum. From right to left : Ps ¼ 6.5, 7.5, 8.5, 9.5,
10.5 mbar.

We note that in absence of saturation, the product Q.Ps
(partial pressure of H2O at the surface), shown in abscissa,
is directly proportional to the H2O column density.
Calculations show that this applies for water partial
pressures typically lower than 0.002 mbar, which approxi-
mately corresponds to a water vapor column density of
20 pr-mm. This is illustrated in Fig. 3b, which shows the
H2O column density as a function of the H2O surface
mixing ratio for a given set of surface pressures, calculated
for a cold temperature profile, for which the saturation
effect is maximum. This profile corresponds to the atmo-
spheric conditions of early northern summer at high
northern latitudes (Ts ¼ 250 K, T(z ¼ 0) ¼ 220 K). As the
water content increases, saturation occurs at lower and
lower altitudes; in this case, saturation takes place directly
at the surface for H2O mixing ratios higher than 3.5 10-3;
the corresponding water vapor column density is then
140 pr-mm.
It should be mentioned that the water vapor column
density is more appropriate than the H2O surface-mixing
ratio to present our results because, to first order, it is the
physical quantity inferred from our analysis. This para-
meter is thus used in the following sections.
In order to determine the surface pressure, we first
considered the measurement of the CO2 band depth. In
particular, the weak 1.44 mm band shows an excellent
linearity up to surface pressures of 10 mbar. This method,
however, has a severe limitation associated to the presence
of water ice signature in some OMEGA spectra, especially
over the polar caps (Langevin et al., 2005a). An example of
such spectra is shown in Fig. 4. In the presence of water ice,
strong signatures appear around 1.5 and 2.0 mm so that the
1.44 mm CO2 band is no more detectable. The water ice
signature appears over the polar caps, but also at locations
where water ice clouds are present. For this reason, we
have chosen to use the surface pressure predictions given
by the EMCD on the basis of the interpolation of the
MOLA data. We have checked that, in the spectra where
the water ice signature is absent, the 1.44 mm CO2 band is
well fitted using the EMCD values of the surface pressure.
Fig. 5 shows the inferred H2O column density for orbit
0037_3 (Ls ¼ 3371), which was used as reference. The
maximum H2O column density, north of Olympus summit,
is about 15 pr-mm, in good agreement with previous
determinations (Smith, 2002, 2004). In this case, saturation
is negligible and the H2O column density is directly related
to the H2O partial pressure at the surface.
4. Results
4.1. The equinox period (Ls ¼ 330–401)
Between Ls ¼ 3301 and 401, i.e. before and after north-
ern spring equinox, the water vapor content is expected to
be rather homogeneous over the Ls range, with latitude
variations ranging from about 3 to 15 pr-mm (Smith, 2002,
2004). Fig. 6 shows the H2O mixing ratio Q and the H2O
column density as a function of latitude and longitude. We
note that the water column density mostly reflects the
topography, and that water saturation is again negligible.
The sampling in latitude and longitude will be hopefully
completed at the end of the extended Mars Express
mission.
4.2. The water sublimation during early northern summer
(Ls ¼ 93–1261)
Figs. 7–12 show the water vapor column density at high
northern latitudes for 6 time steps (Ls ¼ 93–1001,
101–1051, 106–1101, 111–1151, 116–1201, 121–1261). The
geometrical configuration of Mars Express orbits allowed
us to obtain a complete coverage of the northern polar cap
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0.8
0.6
0.4
0.2
reflectanceratio
0.0
1.0 1.5 2.0 2.5
wavelength (µm)
Fig. 4. Examples of reflectance factors in the northern polar region, with
and without the presence of water ice. Data are taken from orbit 902_1.
Black line : lat. above 81N; red line : latitude below 801N. The black
spectrum shows two strong and broad bands at 1.5 and 2.0 mm,
characteristic of water ice. The red spectrum, taken at lower latitudes,
does not show the water ice signatures.
Fig. 5. H2O column density recorded over Olympus Mons (Orbit 0037_3,
Ls ¼ 3371).

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Fig. 7. H2O column density at high northern latitudes for Ls ¼ 93–1001.
The maximum column density is 70 pr-mm.
Fig. 6. A map of the H2O column density for Ls ¼ 330–401 (orbits 6–520),
as a function of latitude and longitude.
Fig. 10. H2O column density at high northern latitudes for
Ls ¼ 111–1151. The maximum column density is 70 pr-mm.

over this time sequence. It can be seen that the water vapor
content increases at the beginning of the sequence and
remains approximately constant over the range
Ls ¼ 100–1151. It may show a second increase for
Ls41201, but the spatial coverage is no more sufficient
for a conclusive statement.
We note that the OMEGA data provide for the first time
the distribution of the water vapor content above the polar
cap, as this region is not observable in the thermal infrared,
because of its low surface temperature. In addition, the
high spatial resolution provided by the OMEGA data
allows for the first time to correlate the atmospheric water
content with the fine structure of the north polar cap, as
shown by the MOLA (version 1/32) data (Fig. 13). It can
be seen that enhancements of the water vapor column
density are closely correlated to altitude changes in the
polar cap. In addition, enhanced water abundance is
clearly visible at the edge of the polar cap, at latitudes
75–801N, especially at longitudes 210–2401E. We note that
this region is not far from the dark longitudinal dunes of
Olympus Planitia where OMEGA detected a large content
of carbon-rich sulfate (80–851N), most likely gypsum
(Langevin et al., 2005a, b). However, the two areas do
not coincide exactly, so there is no clear evidence for a
correlation between the presence of doubly hydrated
sulfate (CaSO4, 2H2O) at the surface and an increased
amount of atmospheric water vapor. We note that the
variability of the water vapor detection from different
orbits covering this region proves that the 2.5 mm sulfate
band does not influence our measurements.
An important comment has to be made about Fig. 7
(Ls ¼ 93–1001). It can be seen that the central blue region
(corresponding to a minimum value of the water vapor
content) seems to be correlated with the maps of small
grains of H2O ice at the surface, obtained by OMEGA for
Ls ¼ 931, (Langevin et al., 2005b, Fig. 2), This suggests
that, in this case, our determination of H2O is most likely
contaminated by the spectrum of these small icy grains,
which have a typical size of about 10–100 mm. As shown in
Fig. 14, the laboratory spectrum of small icy grains
(Schmitt et al., 2003) exhibits a small local maximum
exactly at 2.6 mm (see also Langevin et al., 2005b, Fig. 3).
As a result, the apparent depth of the water vapor band,
measured with respect to this continuum, tends to be
underestimated and the inferred H2O column density is
underestimated accordingly. It can be seen from Fig. 14
that the ratio [2 Â I(2.60 mm)/I(2.55 mm)+I(2.65 mm)] is
about 2% for 10-mm grains and 6% for 100-mm grains.
In the ‘‘snake-like’’ blue region shown in Fig. 7 where
the inferred water vapor content is minimum (about
5–10 pr-mm), the measured H2O line depth is about 2–3%
(Fig. 3a). Correcting this value with the contribution due to
the H2O ice signature would lead to an enrichment of the
H2O water vapor up to about 30–50 pr-mm, As a result,
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Fig. 13. Altimetry map of the northern polar cap (MOLA 32 data).
1.0
0.8
0.6
0.4
0.2
0.0
1.0 1.5 2.0
wavelength (µm)
2.5
reflectancefactor
Fig. 14. Laboratory spectrum of H2O snow for different grain size (from
top to bottom, 10, 100 and 1000 mm). The figure is taken from Schmitt
et al. (2003).

there is no more evidence for a decrease of the water vapor
content above the pole. This is also confirmed in Fig. 8-12,
which show a progressive decrease of the dark-blue ‘‘snake-
like’’ feature and a progressive increase of the H2O column
density over the poles as the small icy grains evaporate
(Langevin et al., 2005b). In conclusion, we interpret the
apparent decrease of the H2O column density over the pole
as due to the presence of small icy grains, which evaporate
progressively, and we see no evidence for a decrease of the
water vapor content over the pole with respect to the
surrounding areas.
5. Discussion
Before comparing our results to previous water vapor
data sets, we need to assess the uncertainty associated to
our measurements.
5.1. Uncertainty analysis
There are several sources of uncertainties in our analysis:
signal-to-noise ratio of the data, non-linearity of the
detector response, water content in the reference spectrum,
modelling uncertainty associated to the atmospheric
parameters.
5.1.1. Intrinsic S/N in an individual spectrum
The 1Às noise level of an individual spectrum is
estimated to 2 DN, while the continuum level, at 2.6 mm,
ranges from about 8000 DN for the brightest spectra,
recorded under maximum insolation (Fig. 1) down to
about ten times less for regions showing the water ice
signature (Fig. 2). The 1Às S/N per individual spectrum
thus ranges from about 400 to 4000 in the 2.6 mm
continuum, inside and outside the polar cap, respectively.
As all spectra are ratioed to a given reference, the 1Às S/N
per individual spectrum ranges from about 300 to 3000.
The depth of the 2.6 mm H2O band is typically 1–6%, so
that the associated uncertainty ranges between 1% and
60% per spectrum. Figs. 7–12 each correspond to an
integration of about 20–40 orbits, of 5–6 sequences each,
with 64–128 longitudinal pixels and more than 200 pixels
along the latitude axis for each orbital sequence. Each
individual pixel in Figs. 7–12 is the average of about ten
OMEGA pixels, and the associated S/N uncertainty on the
line depth thus varies from 0.3% to 20%; the correspond-
ing uncertainty on the water column density is less than
30%. We believe that this number is conservative, because
it corresponds to the combination of all unfavorable
conditions. In fact, at high northern latitudes, the
continuum level is lower because of the water ice signature,
but the spatial coverage is better as most of the orbits
overlap, so that both effects compensate. In order to take
into account the overlap of the orbits the OMEGA spatial
resolution has been reduced to a grid of 0.11 Â 0.11
(OMEGA maximum resolution is $0.011).
5.1.2. Non-linearity of the detector response
It has been shown that the response of the detector,
if plotted as a function of the received signal, exhibits an
oscillation with amplitude of about a percent. The effect is
stronger when the continuum is high. The non-linearity of
the detector response is expected to limit the S/N of our
data in the bright regions, but should not affect the polar
caps areas.
5.1.3. Water content in the reference spectrum
It has been mentioned above that the uncertainty about
the water column density above Olympus Mons is
estimated to be $2 pr-mm. Thus, this source of uncertainty
is dominant for the data set acquired at Ls ¼ 330–401
(Fig. 6), where the water vapor column density is typically
10 pr-mm. On the other hand, this source of error is only a
few percent for the maps obtained at Ls ¼ 93–1261, and is
then a minor contributor to the total uncertainty.
5.1.4. Uncertainties associated to atmospheric modelling
The main parameter in the retrieval of the water vapor
column density is the surface pressure. The Ps values
inferred from the GCM on the basis of the MOLA data
interpolation are believed to be accurate within a few
percent (Forget et al., 1999). However, the spatial
resolution of our surface pressure map has been chosen
in order to take into account one pixel any 10 over the
longitude and 1 pixel over 40 over the latitude, creating a
grid of $0.31 Â 0.31 which strongly depends on the
observation, this choice does not influence our calculations
because of the OMEGA grid of 0.11 already mentioned.
With regard to the temperature dependence, as mentioned
above, calculations show that the measured depth of the
2.6 mm H2O band vary by less than 3% when the
temperature profile varies over the whole range of expected
seasonal variations. The associated uncertainty on the H2O
water content is about twice this value.
In summary, taking into account all possible sources of
errors, we estimate that our error bar on the H2O column
density is about 35% for the Ls ¼ 330–401 period, and is
less than 30% for the Ls ¼ 93–1261 period. We take this
value as our 1Às error bar for the whole OMEGA data set.
Obviously some areas have a lower error bar, either
because they correspond to a strong continuum at 2.6 mm
(absence of water ice) or because they correspond to
multiple orbit overlaps. The quality of the OMEGA data is
also illustrated by the repeatability of the water maps in
Figs. 8–10.
5.2. Comparison with previous data sets
Two global data sets of the water vapor column density,
over a whole seasonal cycle, have been obtained by space
missions: the MAWD-Viking data (Jakosky and Farmer,
1982) and the TES-MGS data (Smith, 2002, 2004). Tables 1
and 2 show a comparison of our results, integrated over the
ARTICLE IN PRESS

longitude, with these two data sets, for the two periods
discussed in this paper.
As shown in Figs. 5–13, our data show strong long-
itudinal variations. This is especially true in Fig. 13 for the
region of enhanced water vapor content corresponding to
Olympia Planitia, but it is also clearly visible in Fig. 6b.
In spite of this limitation, Tables 1 and 2 show that, while
the results are in good general agreement for the equinox
region, our results are significantly lower than previous
analyses in the case of the northern polar cap sublimation.
It should be mentioned, however, that, while the first TES
results indicated maxima of 60–75 pr-mm (Smith, 2002), in
the latest TES study (Smith, 2004) no absolute value is
given for the maximum water content at Ls ¼ 100–1201;
the only indication given on the map is that the maximum
value is higher than 40 pr-mm, in good agreement with our
results. Finally, we also note that our last map, which
unfortunately shows an incomplete spatial coverage, seems
to indicate a local water vapor enhancement with values
above 60 pr-mm.
In conclusion, the present analysis of the OMEGA data
set shows that a retrieval of the water vapor content in the
atmosphere is possible. However, above the north pole, a
precise retrieval of the water vapor content is not possible
because of the presence of small-size grains at the surface.
The main conclusions of this paper can be summarized
as follows:
High-resolution water vapor maps above of the north-
ern polar region have been retrieved for the first time, at
the time of water vapor sublimation. However, above
the pole itself, no precise measurement is possible due to
the presence of small icy grains.
An enhancement of the water vapor content appears at
latitudes 75–801N and longitude 210–2401E. We note
that this region is not far from the area where carbon-
rich hydrated sulfates have been identified in large
amounts, however there is no evidence for a clear
correlation.
The water column densities derived at high northern
latitude during early northern summer appear lower
than the results inferred from previous studies.
In the future, this analysis will be extended to the whole
OMEGA data set and will hopefully cover the whole
seasonal water cycle at the end of the Mars Express
extended mission.
References
Bibring, J.-P., Erard, S., Gondet, B., et al., 1991. Topography of the
martian tropical regions with ISM. Plan Space Sci. 39, 225–236.
Bibring, J.-P., Soufflot, A., Berthe´ , M., et al., 2004. OMEGA:
Observatoire pour la Mine´ ralogie, l’Eau, les Glaces et l’Activite´ .
ESA-SP 1240, 37–49.
Chicarro, A., Martin, P., Trautner, R., 2004. The Mars Express mission:
an overview. ESA-SP 1240, 3–13.
Clancy, R.T., Grossman, A.W., Muhleman, D.O., 1992. Mapping Mars
water vapor with the Very Large Array. Icarus 100, 48–59.
Clancy, R.T., Grossman, A.W., Wolff, M.J., James, P.B., Rudy, D.J.,
Billawala, Y.N., Sandor, B.J., Lee, S.W., Muhleman, D.O., 1996.
Water vapor saturation at low altitudes around Mars aphelion: a key
to Mars climate? Icarus 122, 36–62.
Conrath, B., Curran, J., Hanel, R., Kunde, V., Maguire, W., Pearl, J.,
Pirraglia, J., Welker, J., 1973. Atmospheric and surface properties of
Mars obtained by spectroscopy on Mariner 9. J. Geophys. Res. 78,
4267–4278.
Encrenaz, T., Lellouch, E., Rosenqvist, J., et al., 1991. The atmospheric
composition of Mars: ISM and ground-based observational data. Ann.
Geophys. 9, 797–803.
Encrenaz, T., Lellouch, E., Cernicharo, J., Paubert, G., Gulkis, S., 1995.
A tentative detection of the 183-GHz water vapor line in the martian
atmosphere: constraints upon the H2O abundance and vertical
distribution. Icarus 113, 110–118.
Encrenaz, T., Lellouch, E., paubert, G., Gulkis, S., 2001. The water vapor
vertical distribution on Mars from millimeter transitions of HDO and
H2
18
O. Planet. Space Sci. 49, 731–741.
Encrenaz, T., Melchiorri, R., Fouchet, T.etal., 2005. A mapping of
martian water sublimation during early northern summer using
OMEGA/Mars Express. Astron. Astrophys. 441, L9–L12.
Farmer, C.B., Davis, D.W., Holland, A.L., et al., 1977. Mars: Water
vapor observations from the Viking orbiters. J. Geophys. Res. 82,
4225–4248.
Fedorova, A.A., Rodin, A.V., Baklanova, I.V., 2004. MAWD observa-
tions revisited: seasonal behavior of water vapor in the martian
atmosphere. Icarus 171, 54–67.
Forget, F., Hourdin, F., Fournier, R., et al., 1999. Improved general
circulation models of the Martian atmosphere from the surface to
above 80 km. J. Geophys. Res. 104, 24155–24176.
Jacquinet-Husson, N., et al., 1999. The 1997 spectroscopic GEISA
databank. J. Quant. Spectrosc. Rad. Transfer 62, 205–254.
Jakosky, B.M., Farmer, C.B., 1982. The seasonal and global behavior of
water vapor in the Mars atmosphere: complete global results of the
Viking atmospheric water detector experiment. J. Geophys. Res. 87,
2999–3019.
Langevin, Y., Poulet, F., Bibring, J.-P., Schmitt, B., Doute´ , S., Gondet,
B., 2005a. Summer evolution of the north polar cap of Mars as
observed by OMEGA/Mars Express. Science 307, 1581–1584.
ARTICLE IN PRESS
Table 1
Comparison of water vapour data sets, Ls ¼ 330–401
Latitude MAWD-Viking TES-MGS This work
501N 5À10 5À10 5À10
301N 10 10–15 10À15
0 10 10 10À15
301S 5À10 5À10 10
601S 1À5 5 5À10
801S o3 o5
Table 2
Comparison of water vapour data sets, Ls ¼ 93–1261
Latitude MAWD-
Viking
TES-MGS This work
701N 70 60 30–50
801N 80 75 40–70
901N 30 (without the
granular icy region)

Langevin, Y., Poulet, F., Bibring, J.-P., Gondet, B., 2005b. Sulfates in the
north polar region of Mars detected by OMEGA/Mars Express.
Science 307, 1584–1586.
Lellouch, E., Encrenaz, T., de Graauw, T., et al., 2000. The 2.4-45 micron
spectrum of Mars observed with the Infrared Space Observatory.
Planet. Space Sci. 48, 1393–1405.
Montmessin, F., Forget, F., Rannou, P., Cabanne, M., Haberle, R.M.,
2004. Origin and role of water ice clouds in the Martian water cycle as
inferred from a general circulation model. J. Geophys. Res. 109, E10
CiteID E10004.
Mustard, J.F., Poulet, F., Gendrin, A., et al., 2005. Olivine and pyroxene
diversity in the crust of Mars. Science 307, 1594–1597.
Pollack, J.B., Dalton, J.B., Grinspoon, D., et al., 1993. Near-infrared light
from Venus’ nightside—a spectroscopic analysis. Icarus 103, 1–42.
Schmitt, B., Doute´ , S., Erard, S., Langevin, Y., Bibring, J.-P., 2003.
Discrimination between solid, liquid and gaseous H2O, and assessment
of H2O temperature by spectroscopy using the OMEGA imaging
spectrometer (Mars Express). Communication presented at the work-
shop )Planet Mars*, April 28—May 9, Les Houches, France.
Smith, M.D., 2002. The annual cycle of water vapor on Mars as observed
by the Thermal Emission Spectrometer. J. Geophys. Res. 107, 1
doi: 10.1029/2001/JE001522, E11, 5115.
Smith, M., 2004. Interannual variability in TES atmospheric observations
of Mars during 1999–2003. Icarus 167, 148–165.
Sprague, A., Hunten, D.M., Hill, R.E., Rizk, B., Wells, W.K., 1996. Mars
water vapor: 1988–1995. J. Geophys. Res. 101, 23229–23254.
Sprague, A., Hunten, D.M., Doose, L.R., Hill, R.E., 2003. Mars
atmospheric water vapor abundance: 1996–1997. Icarus 163, 88–101.
Titov, D.V., Markiewicz, W.J., Thomas, N., Keller, H.U., Sablotny,
R.M., Tomasko, M.G., Lemmon, M.T., Smith, P.H., 1999. Measure-
ments of the atmospheric water vapor on Mars by the Imager for Mars
Pathﬁnder. J. Geophys. Res. 104, 9019–9026.
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