Icarus 178 (2005) 487–492
www.elsevier.com/locate/icarus
A sensitive search for SO2 in the martian atmosphere:
Implications for seepage and origin of methane
Vladimir A. Krasnopolsky ∗,1
Department of Physics, Catholic University of America, Washington, DC 20064, USA
Received 2 December 2004; revised 29 April 2005
Available online 1 July 2005
Abstract
Mars was observed near the peak of the strongest SO2 band at 1364–1373 cm−1 with resolving power of 77,000 using the Texas Echelon
Cross Echelle Spectrograph on the NASA Infrared Telescope Facility. The observation covered the Tharsis volcano region which may be
preferable to search for SO2 . The spectrum shows absorption lines of three CO2 isotopomers and three H2 O isotopomers. The water vapor
abundance derived from the HDO lines assuming D/H = 5.5 times the terrestrial value is 12 ± 1.0 pr. µm, in agreement with the simultaneous
MGS/TES observations of 14 pr. µm at the latitudes (50◦ S to 10◦ N) of our observation. Summing of spectral intervals at the expected
positions of sixteen SO2 lines puts a 2σ upper limit on SO2 of 1 ppb. SO2 may be emitted into the martian atmosphere by seepage and is
removed by three-body reactions with OH and O. The SO2 lifetime, 2 years, is longer than the global mixing time 0.5 year, so SO2 should
be rather uniformly distributed across Mars. Seepage of SO2 is less than 15,000 tons per year on Mars which is smaller than the volcanic
production of SO2 on the Earth by a factor of 700. Because CH4 /SO2 is typically 10−4 –10−3 in volcanic gases on the Earth, our results
show seepage is unlikely to be the source of the recently discovered methane on Mars and therefore strengthen its biogenic origin.
 2005 Elsevier Inc. All rights reserved.
Keywords: Mars, atmosphere; Atmospheres, composition; Spectroscopy; Infrared observations; Exobiology
1. Introduction
Other than water vapor and CO2 , which are typical of
the martian atmospheric environment, sulfur dioxide (SO2 )
is the most abundant species in terrestrial volcanic gases
and constitutes a few per cent of total gas release (Holland,
1978). However, the latest traces of active martian volcanism are 2 to ∼100 million years old (Hartmann and Berman,
2000; Sakimoto et al., 2003; Neukum et al., 2004), and the
thermal emission imaging system which was specially designed to search for hot spots on Mars from Mars Odyssey
* Correspondence address: 6100 Westchester Park #911, College Park,
MD 20740, USA.
E-mail address: [email protected].
1 Visiting Astronomer at the Infrared Telescope Facility, which is operated by the University of Hawaii under Cooperative Agreement No. NCC
5-538 with the National Aeronautics and Space Administration, Office of
Space Science, Planetary Astronomy Program.
0019-1035/$ – see front matter  2005 Elsevier Inc. All rights reserved.
doi:10.1016/j.icarus.2005.05.006
orbiter, has not detected any hot spots (Christensen, 2003).
Yet weak seepage of the volcanic and geo- or hydrothermal
gases is still possible, and photochemical models for those
hypothetical sources of SO2 , H2 S, and CH4 have been calculated (Wong et al., 2003).
SO2 may also be extracted by the solar UV photons from
sulfates on the martian surface or airborne dust. Sulfates
have not been detected on Mars by the thermal emission
spectrometer on Mars Global Surveyor (McSween et al.,
2003; Wyatt et al., 2003). However, both the α-particle X-ray
spectrometer and the miniature thermal emission spectrometer on the Opportunity rover revealed significant (15 to
35% by volume in some locations) amounts of magnesium
and calcium sulfates (Rieder et al., 2004; Christensen et al.,
2004).
The recent discovery of methane on Mars (Krasnopolsky
et al., 2004; Formisano et al., 2004; Mumma et al., in preparation) triggered the discussion of its biogenic and abiogenic
488
V.A. Krasnopolsky / Icarus 178 (2005) 487–492
sources. SO2 may be an effective tracer of the current outgassing on Mars.
An upper limit of 100 parts per billion (ppb) was established for SO2 in the martian atmosphere based on the
absence of the absorption band at 7.3 µm in the Mariner 9
infrared spectra (Maguire, 1977). Later this limit was improved to 30 ppb using a microwave observation at 1.4 mm
(Encrenaz et al., 1991). However, even an SO2 abundance
below this limit may affect martian photochemistry, and
some of the models by Krasnopolsky (1993, hereafter Paper I) were calculated assuming a SO2 mixing ratio of
10 ppb.
2. Observation
To search for SO2 on Mars, we also used the strongest
band at 7.3 µm. The observation was made on 18 June 2003,
when Mars had a diameter of 14.7 arcsec, phase (Sun–Mars–
Earth) angle 40◦ , and geocentric velocity −11.0 km s−1 .
We used the NASA Infrared Telescope Facility on Mauna
Kea, Hawaii. A comparatively low atmospheric pressure of
0.6 bar and dry atmosphere with a typical water abundance
of 2 pr. mm above Mauna Kea (elevation 4.2 km) are essential for the observation near 7.3 µm where both H2 O and
CH4 telluric lines are strong and the atmospheric opacity is
significant even between the telluric lines. The observation
is essentially impossible at low-altitude observatories.
We used TEXES, the Texas Echelon Cross Echelle Spectrograph (Lacy et al., 2002), in its high resolution mode.
The instrument slit, 1.5 × 7 arcsec2 , was parallel to the central meridian and between the subsolar point and the central meridian. Its position corresponded to local time near
13:30 when the thermal contrast between the surface and
the atmosphere is close to maximum. The observation was
at subsolar longitude LS = 205◦ (southern spring), when
the heliocentric distance was near the minimum (1.42 and
1.38 AU, respectively), and the slit covered latitudes from
50◦ S to 10◦ N. The longitude changed from 104 to 164◦ W
during four hours of the observation. Our observation covered the Tharsis volcano region which may be preferable for
a search for SO2 . We observed Mars, a flat field (an outer ambient temperature black body), the Moon, two standard stars,
and sky foreground 30 arcsec north of all targets. The total
on-Mars exposure was 31 min of four hours of the observing
time. Because Mars is such a bright source, we increased the
amount of time spent observing the blackbody to insure the
final result was not limited by photon noise from the blackbody.
The instrument spectral coverage near the peak of the
SO2 band at 1362 cm−1 is 9 cm−1 . Choosing a spectral range
should be done carefully to cover the strongest lines and
to minimize the contamination by telluric lines. We chose
1364 to 1373 cm−1 which includes two narrow opaque intervals at 1368.2–1369.0 and 1372.1–1373.2 cm−1 centered
at the strong H2 O lines. For example, the broad intervals at
1360.5–1363.7 and 1373.0–1375.7 adjacent to our spectral
interval are almost completely opaque. The chosen spectral
range covers many strong SO2 lines as can be seen in Fig. 1.
Fig. 1. Spectrum of the martian atmosphere. The spectrum consists of the martian lines of three isotopomers of CO2 and three isotopomers of H2 O as well as
some residuals of the telluric CH4 and H2 O lines. The spectrum is shifted at rest position (corrected for the Doppler shift). Positions of seventeen SO2 lines
chosen for extraction are shown by bars whose sizes reflect the line strengths. Two lines at 1370.32 cm−1 coincide, and their sum is given.
Search for SO2 on Mars
3. Results
The data were processed using typical TEXES reduction
software. The software differences nod pairs, flatfields, corrects for spikes, produces a wavenumber scale based on atmospheric features, and makes a significant reduction of the
telluric absorption lines and emissions from the telescope
and optics by subtraction of the sky foreground. Even better
compensation for the telluric lines may be achieved using
the Moon spectra. However, this results in some increase of
the noise, and we do not use this possibility. The data were
resampled and spectral regions with usable data in two orders were combined. To search for SO2 , we summed all our
data over time and spatial position along the slit. The resulting spectrum is shown in Fig. 1. This spectrum is a ratio of
the Mars minus sky to the blackbody minus sky spectra. This
ratio corrects for variations in the pixel sensitivity (flat field).
Telluric absorption lines are significantly reduced in this
spectrum. Small variations of the continuum were corrected
using polynomial fitting. The mean continuum brightness is
25 erg (cm2 s sr cm−1 )−1 and corresponds to the surface temperature of 276 K. This temperature is close to that expected
for the season, latitudes, and local time (13:30) of our observation from the Viking IRTM data (Martin, 1981). The
temperature is high because the observation was near the
daytime temperature maximum and Mars was rather close
to perihelion.
The spectrum shows lines from three isotopes of CO2
and three isotopes of H2 O in the martian atmosphere and
some residuals of telluric CH4 and H2 O lines. Comparing wavenumbers of ten CO18 O lines of the strong 10001–
00001 band in our spectrum with their wavenumbers in
the HITRAN 2004 spectroscopic database, we found a systematic difference of (56.9 ± 0.5) × 10−3 cm−1 , of which
0.050 cm−1 is the Doppler shift and the remaining (6.9 ±
0.5) × 10−3 cm−1 represents a systematic correction to the
wavenumber. Therefore the wavenumber scale in Fig. 1 has
been corrected by 0.057 cm−1 , and the spectrum is at rest
position.
The martian weak lines are very narrow, and their shape
in the observed spectrum reflects the instrument response
function. We fitted some weak lines by the Gaussian and
Lorentzian profiles and found that the Gaussian fit is much
better. A mean dispersion of this fit is equal to 3.3 ± 0.1 subpixel intervals of d = 0.00228 cm−1 . (The instrument pixel
size is twice the subpixel value. The spectral pixel size was
reduced by a factor of 2 to avoid blurring when data from
two diffraction orders were added together. The wavenumber sampling slightly varies from the beginning to the end
of the spectrum which is typical of grating spectrographs.)
This dispersion corresponds to a spectral resolution element
δν = 0.0177 cm−1 (full width at half maximum) and the instrument spectral resolving power ν/δν = 77,000.
For our analysis of the martian lines we adopt their
strengths at a temperature at a level of half atmospheric
pressure. This approach is similar to the Curtis–Godson
489
approximation for the strong collisionally broadened lines
(Chamberlain and Hunten, 1987) and provides an uncertainty of 4% for martian lines (Krasnopolsky et al., 1997).
This uncertainty is acceptable for our goals. According to
the MGS/TES simultaneous observations (Smith, 2004), the
mean temperature at the level of half atmospheric pressure
for latitudes of 50◦ S to 10◦ N was equal to 223 K.
The martian lines are very narrow; therefore we did not
fit the observed lines by a synthetic spectrum and examined their equivalent widths instead. Equivalent widths of
six weak CO2 lines were used to determine an effective
CO2 abundance of (2.0 ± 0.2) × 1023 cm−2 . (Effective abundance is the derived abundance before correcting for emission by the absorbing molecule.) The globally and annually
mean atmospheric pressure is 6.1 mbar on Mars. Using the
Viking lander measurements (Tillman, 1988), the mean pressure is equal to 5.8 mbar at the season of our observation
(LS = 205◦ ). Using the MOLA data, we find that the mean
elevation of the observed region is 3.0 km and the mean
pressure is 4.5 mbar at the observed region. Then the actual CO2 abundance is 1.7 × 1023 cm−2 , and a correction
for the mean air mass of 1.15 in the martian atmosphere results in 2.0 × 1023 cm−2 . The weak CO2 lines correspond
to absorption of the surface radiation with T = 276 K and
emission with the mean temperature of 223 K. The emitted radiation is weaker than the absorbed radiation by a
factor of e1.439ν(1/223−1/276) = 5.5, where ν is the wavenumber. Therefore the effective CO2 abundance derived from the
weak CO2 lines should be equal to the true slant CO2 abundance times a factor of 1 − 5.5−1 = 0.82, that is, 1.64 ×
1023 cm−2 . The measured effective CO2 abundance slightly
exceeds the calculated value.
The HDO lines at 1365.301, 1367.574, and 1372.015
cm−1 give a water vapor abundance of 12 ± 1 pr. µm. D/H
on Mars is 5.5 times the terrestrial value (Owen et al., 1988;
Krasnopolsky et al., 1997), and this ratio has been applied.
Saturation of water vapor occurs at h 30 km at the season
and latitudes of our observation and water vapor may be considered as uniformly mixed in the lower atmosphere. The derived water vapor abundance is close to that observed simultaneously with the MGS/TES, which is equal to 14 pr. µm
for the latitudes of our observation (Smith, 2004).
To search for SO2 in the spectrum, we examine the positions of seventeen lines which are strong, not contaminated
by other lines, and have low noise. (Noise is variable in our
spectrum.) We compared high-precision heterodyne measurements of frequencies for some of the SO2 lines of our
interest (Vanek et al., 1990; Flaud et al., 1993) with those
given in HITRAN 2004 and concluded that an uncertainty
of the chosen SO2 line positions is 0.0003 cm−1 . This uncertainty is much smaller than our resolution element, and
the line positions from HITRAN 2004 may be used with a
good confidence. The chosen seventeen SO2 lines and their
strengths are shown in Fig. 1.
Thirteen subpixels centered at the expected position of
a chosen SO2 line are used for the extraction of SO2 . Four
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V.A. Krasnopolsky / Icarus 178 (2005) 487–492
Fig. 2. Top: 21 points of the spectrum centered at two coinciding SO2 lines
at 1370.3209 and 1370.3214 cm−1 . Thin line is a third-order polynomial
fitting to four left and four right points. This fit is a continuum model. Bottom: difference between the spectrum and the continuum is compared with
the SO2 absorption line (thin line) calculated for a mixing ratio of 1 ppb.
subpixels to the left and to the right of these 13 subpixels
are approximated by a third degree polynomial to model a
continuum near the line. (Spectra around some lines have
inflection points and therefore need third or higher degree
polynomials for approximation. We use the third degree
polynomials, and just two parameters, intensity and slope,
are actually taken from the four points at the each side of the
line.)
Two strong SO2 lines at 1370.3209 and 1370.3214 cm−1
blend together within our subpixel size, and a sum of their
strengths significantly exceeds the strengths of the other
lines (Fig. 1). 13 central and 8 outer points for this line and
the modeled continuum are shown in Fig. 2 (top). A difference between the spectrum and the continuum is shown in
Fig. 2 (bottom) where it is compared with a SO2 line shape
calculated for a SO2 mixing ratio of 1 ppb. The SO2 line
equivalent width is equal to the CO2 effective abundance
(2 × 1023 cm−2 ) times the adopted SO2 mixing ratio times
the line strength. Then this equivalent width is scaled to the
instrument line shape which is a Gaussian with FWHM of
the instrument resolution. This approach works if both the
CO2 and SO2 lines are weak and SO2 /CO2 is constant with
height. (This will be discussed below.) This analysis avoids
radiative transfer modeling which involves thermal properties of the atmosphere and the surface.
The spectrum in Fig. 2 is on the wing of the strong CO18 O
line, and the continuum fitting by the third degree polynomial may be not perfect for this spectrum. Therefore the
difference between the spectrum and the continuum in the
lower panel of Fig. 2 is actually a sum of the noise, a possible
error in the continuum fitting, and a possible SO2 absorption.
The minimum seen in the figure is displaced by 0.009 cm−1
which significantly exceeds the wavenumber uncertainty of
Fig. 3. Sum of spectral intervals centered at the expected positions of sixteen
SO2 lines and corrected for their continua (see text). Error bars show standard deviations of the summed points. Each subpixel is d = 0.00228 cm−1 .
S = 1.0 × 10−18 cm is the sum of the sixteen line strengths. The Gaussian
has a width of the instrument spectral resolution (0.0177 cm−1 ) and corresponds to the SO2 mixing ratio of 1 ppb in the martian atmosphere.
0.0005 cm−1 in our spectrum. Therefore, the spectrum in
Fig. 2 does not show the SO2 absorption. To improve the
confidence of our search, differences between the spectrum
and the modeled continuum are summed up for all chosen
lines and shown in Fig. 3. The line strengths are equal to
(4–8) × 10−20 cm at 223 K, and the summed spectrum was
multiplied by the subpixel size and divided by the sum of the
line strengths, S = 1.0×10−18 cm, to give an effective abundance of SO2 . Evidently no absorption has been detected.
To obtain an upper limit to the SO2 abundance, we need
to estimate the uncertainty in the summed spectrum in Fig. 3.
Two methods have been applied. The first method is to estimate the noise in the spectrum and multiply it by square root
of the number of the lines. Four spectrally clean regions at
1365.32–1365.41, 1366.70–1366.90, 1367.36–1367.52, and
1371.06–1371.29 cm−1 were taken to estimate the noise:
(yi − y0i )2 1/2
.
N=
n/2 − k
Here N is the noise, yi is the spectral point, y0i is the point
from linear (k = 2) or quadratic (k = 3) fit to the spectrum,
and n/2 is the true pixel number in the chosen interval. Then
the mean noise from these four regions is N = 0.0012, the
noise in Fig. 3 is 161/2 N d/S = 1.1 × 1013 cm−2 per subpixel, and a two-sigma upper limit for SO2 is 2 × 3.3 ×
(2π)1/2 × 1.1 × 1013 = 1.8 × 1014 cm−2 . (3.3 × (2π)1/2 is
the effective number of subpixels in the Gaussian with the
dispersion of 3.3 subpixels.)
Another method is to use a standard deviation for the
points in the 16 summed spectra. It is similar to the derived noise and also results in the two-sigma upper limit
of 1.8 × 1014 cm−2 . 5% of the signal is beyond of the
thirteen central subpixels, and the corrected upper limit is
1.9 × 1014 cm−2 . This limit divided by the effective CO2
abundance of 2 × 1023 cm−2 corresponds to a two-sigma up-
Search for SO2 on Mars
per limit to the SO2 mixing ratio of 1 ppb. This limit is much
more restrictive than the currently existing limit of 30 ppb
(Encrenaz et al., 1991).
4. Photochemical loss of SO2
A possible photochemical impact of SO2 was considered
in Paper I where the calculated indirect photolysis of O2 in
the reactions
SO + HO2 → SO2 + OH,
SO + O2 → SO2 + O,
S + O2 → SO + O,
equaled the direct photolysis of O2 for the SO2 mixing ratio
of 10 ppb. With the upper limit established here of 1 ppb,
SO2 cannot significantly affect martian photochemistry and
the photochemical effect of SO2 may be neglected.
The derived upper limit may be converted to a limit on
production of SO2 if we calculate its photochemical loss.
Many of the reactions with SO2 considered in Paper I (see
also Wong et al., 2003) result in recycling of SO2 with no
net loss. For example, SO reacts with O, O2 , O3 , and HO2
and returns SO2 ; therefore photolysis is not a sink of SO2 .
Photochemical loss of SO2 is provided by the reactions
SO2 + O + M → SO3 + M;
k1 = 2.6 × 10−33 (T /300)3.6 cm6 s−1
SO2 + OH + M → HSO3 + M;
k2 = 6 × 10−31 (300/T )3.3 cm6 s−1 .
Reaction 2 is followed by
HSO3 + O2 → SO3 + HO2 ,
and SO3 quickly absorbs water to form sulfuric acid H2 SO4 .
Sulfuric acid either reacts with dust or precipitates to the surface and reacts with the rocks. The reaction rate coefficients
are taken from Sander et al. (2003) where they are given for
air and are doubled to account for the higher efficiency of
CO2 as a third body.
SO2 is uniformly mixed in the atmosphere up to 30 km in
Paper I and up to 60 km in Fig. 2 from Wong et al. (2003).
The difference is mainly due to the extremely large abundance of SO2 , 100 ppm, in Wong et al. (2003) which screens
SO2 from photolysis in the lower atmosphere. The model in
Paper I corresponds to the optically thin conditions which
are applicable to our upper limit. Therefore we reduce the
SO2 vertical profile from Paper I by a factor of 10 to fit our
upper limit.
Using densities of O, OH, and CO2 and temperature profiles from model 3 in Krasnopolsky (1995) and the adjusted
model of Nair et al. (1994), we calculate the total loss of
SO2 in both models. The reaction with OH is more effective
in removal of SO2 than the reaction with O by a factor of
6 in both models. The models result in a lifetime of SO2 on
Mars of 2.4 and 1.5 years, respectively. We adopt the lifetime
of 2 years. It is longer than the time for global atmospheric
491
mixing (0.5 year, see Krasnopolsky et al., 2004), and SO2
should generally be uniformly mixed in altitude up to 30 km
and across the planet. (SO2 cannot condense anywhere on
Mars including the polar caps where the temperature cannot fall below the CO2 condensation temperature of 145 K
at 6 mbar.) However, the difference in the times is not large,
and some low variations in SO2 are not ruled out. Loss and
production of SO2 are balanced, and the calculated upper
limit to these processes is 17,000 t y−1 (tons per year).
5. Discussion
Our upper limit of 1 ppb to SO2 in the martian atmosphere
and the comparatively long lifetime of SO2 do not support
the existence of extended regions with a well-developed photochemistry where the local densities of SO2 may exceed the
measured limit by a few orders of magnitude (Wong et al.,
2003). Furthermore, if some vents exist then their products
are quickly blown off, restricted to a few hundred meters in
altitude, and do not form a well-developed photochemistry
up to 100 km.
Our upper limit of 17 ktons per year to the SO2 production may be compared with the production of 10 megatons
of SO2 annually by volcanoes on Earth (Yung and DeMore,
1999). Currently there is no active volcanism on Mars, and
traces of the latest active volcanism are 2 to 100 million
years old (Hartmann and Berman, 2000; Sakimoto et al.,
2003; Neukum et al., 2004). Delivery of SO2 by cometary
impacts may be calculated using a total mass influx from
comets to Mars, 1000 t y−1 (Krasnopolsky et al., 2004), and
the abundance of sulfur in comet Halley (Jessberger and
Kissel, 1991). The calculated rate is 70 t y−1 of SO2 . Delivery of SO2 by meteorites and interplanetary dust may
be calculated by scaling the delivery of carbon, 1200 t y−1
(Flynn, 1996), to the S/C ratio in meteorites, 0.05 (Anders
and Grevesse, 1989). This results in 330 t y−1 of SO2 .
Therefore, the external sources are weak and negligible.
However, delivery of volcanic and other gases from the interior into the martian atmosphere may be done by seepage
with no eruption of lava. Evidence for recent (less than 106
years old) groundwater seepage comes from MGS images of
Mars that show features such as gullies (Mellon and Phillips,
2001). Therefore, our limit refers to seepage of SO2 and the
solar UV weathering of sulfates on Mars. Our observation
shows that the production of SO2 from both groundwater
and volcanic seepage of SO2 on Mars is weaker than the
production of SO2 by the terrestrial volcanoes by a factor of
more than 600. Mars may be sulfur-rich but much of the sulfur may be locked up in the core as FeS. There are no plate
tectonics on Mars, and this also restricts the delivery of sulfur and other gases into the atmosphere.
Our upper limit is also relevant to the origin of methane
on Mars which was recently detected by three independent
teams (Krasnopolsky et al., 2004; Formisano et al., 2004;
Mumma et al., in preparation). The abundance of SO2 in
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V.A. Krasnopolsky / Icarus 178 (2005) 487–492
terrestrial volcanic gases is a few percent (see above) while
the abundance of CH4 is a few ppm (Etiope and Klusman,
2002). Scaling these values and assuming a similar CH4 /SO2
ratio, we find a limit to seepage of volcanic methane on Mars
of ≈1 t y−1 . This limit is more than two orders of magnitude
smaller than the total production of methane on Mars. Seepage of volcanic methane is a minor source of atmospheric
methane even if CH4 /SO2 on Mars exceeds the terrestrial ratio because of a significant reduction of sulfur in the martian
crust and magma.
The lifetime of methane in the martian atmosphere is
340 years, therefore methane from volcanic eruptions cannot exist in the present atmosphere of Mars. Delivery of
methane by comets, meteorites, and interplanetary dust is
a few percent of the required production. THEMIS at Mars
Odyssey orbiter has not revealed any hot spots on Mars, and
this restricts but does not completely rule out the hydrothermal production of methane. Krasnopolsky et al. (2004) considered these arguments and concluded that production of
methane by methanogenic microbes is a plausible explanation for the detected methane. Now our restriction to seepage of methane strengthens the biogenic source of martian
methane.
Acknowledgments
I am grateful to T.K. Greathouse, M.J. Richter, and J.H.
Lacy who made the observation and the initial data processing, and to M.D. Smith who gave the data of simultaneous
MGS/TES observations of temperature profiles, water vapor,
and dust and ice aerosol. This work was supported by the
NASA Planetary Astronomy Program.
References
Anders, E., Grevesse, N., 1989. Abundances of the elements: Meteoritic and
solar. Geochim. Cosmochim. Acta 53, 197–214.
Chamberlain, J.W., Hunten, D.M., 1987. Theory of Planetary Atmospheres.
Academic Press, Orlando, FL.
Christensen, P.R., 2003. Mars as seen from the 2001 Mars Odyssey thermal
emission imaging system experiment. In: Eos Trans. AGU, Fall Meeting
Suppl. 84 (46). Abstract P21A-02.
Christensen, P.R., 26 colleagues, 2004. Mineralogy at Meridiani Planum
from the Mini-TES experiment on the Opportunity rover. Science 306,
1733–1739.
Encrenaz, T., Lellouch, E., Rosenqvist, J., Drossart, P., Combes, M., Billebaud, F., de Pater, I., Gulkis, S., Maillard, J.P., Paubert, G., 1991. The
atmospheric composition of Mars: ISM and ground-based observational
data. Ann. Geophys. 9, 797–803.
Etiope, G., Klusman, R.W., 2002. Geologic emissions of methane to the
atmosphere. Chemosphere 49, 777–789.
Flaud, J.M., Perrin, A., Salah, L.M., Lafferty, W.J., Guelachvili, G., 1993.
A reanalysis of the (010), (020), and (001) rotational levels of 32 S16 O2 .
J. Mol. Spectrosc. 160, 272–278.
Flynn, G.J., 1996. The delivery of organic matter from asteroids and comets
to the early surface of Mars. Earth Moon Planets 72, 469–474.
Formisano, V., Atreya, S., Encrenaz, T., Ignatiev, N., Giuranna, M., 2004.
Detection of methane in the atmosphere of Mars. Science 306, 1758–
1761.
Hartmann, W.K., Berman, D.C., 2000. Elysium Planitia lava flows: Crater
count chronology and geological implications. J. Geophys. Res. 105,
15011–15025.
Holland, H.D., 1978. The Chemistry of the Atmosphere and Oceans. Wiley,
New York.
Jessberger, E.K., Kissel, J., 1991. Chemical properties of cometary dust and
a note on carbon isotopes. In: Newburn, R., Neuegebauer, M., Rahe, J.
(Eds.), Comets in the Post-Halley Era. Kluwer Academic, Dordrecht,
pp. 1075–1092.
Krasnopolsky, V.A., 1993. Photochemistry of the martian atmosphere
(mean conditions). Icarus 101, 313–332.
Krasnopolsky, V.A., 1995. Uniqueness of a solution of a steady state photochemical problem: Applications to Mars. J. Geophys. Res. 100, 3263–
3276.
Krasnopolsky, V.A., Bjoraker, G.L., Mumma, M.J., Jennings, D.E., 1997.
High-resolution spectroscopy of Mars at 3.7 and 8 µm: A sensitive
search for H2 O2 , H2 CO, HCl, and CH4 , and detection of HDO. J. Geophys. Res. 102, 6525–6534.
Krasnopolsky, V.A., Maillard, J.P., Owen, T.C., 2004. Detection of methane
in the martian atmosphere: Evidence for life? Icarus 172, 537–547.
Lacy, J.H., Richter, M.J., Greathouse, T.K., Jaffe, D.T., Zhu, Q., 2002.
TEXES: A sensitive high-resolution grating spectrograph for the midinfrared. Publ. Astron. Soc. Pac. 114, 153–168.
Maguire, W.C., 1977. Martian isotopic ratios and upper limits for possible minor constituents as derived from Mariner 9 infrared spectrometer
data. Icarus 32, 85–97.
Martin, T.Z., 1981. Mean thermal and albedo behavior of the Mars surface
and atmosphere over a martian year. Icarus 45, 427–446.
McSween, H.Y., Grove, T.L., Wyatt, M.B., 2003. Constraints on the composition and petrogenesis of the martian crust. J. Geophys. Res. 108 (E12),
5135, doi:10.1029/2003JE002175.
Mellon, M.T., Phillips, R.J., 2001. Recent gullies on Mars and the source of
liquid water. J. Geophys. Res. 106, 23165–23179.
Nair, H., Allen, M., Anbar, A.D., Yung, Y.L., Clancy, R.T., 1994. A photochemical model of the martian atmosphere. Icarus 111, 124–150.
Neukum, G., 43 colleagues, 2004. Recent and episodic volcanic and glacial
activity on Mars revealed by the High Resolution Stereo Camera. Nature 432, 971–979.
Owen, T., Maillard, J.P., de Bergh, C., Lutz, B.L., 1988. Deuterium on Mars:
The abundance of HDO and the value of D/H. Science 240, 1767–1771.
Rieder, R., 14 colleagues, 2004. Chemistry of rocks and soils at Meridiani
Planum from the alpha particle X-ray spectrometer. Science 306, 1746–
1749.
Sakimoto, S.E.H., Gregg, T.K.P., Hughes, S.S., Chadwick, J., 2003. Reassessing plains-style volcanism on Mars. In: Sixth International Conference on Mars, Pasadena. Abstract #3197 [CD-ROM].
Sander, S.P., 10 colleagues, 2003. Chemical Kinetics and Photochemical
Data for Use in Atmospheric Studies. Evaluation Number 14. JPL Publication 02-25.
Smith, M.D., 2004. Interannual variability in TES atmospheric observations
of Mars during 1999–2003. Icarus 167, 148–165.
Tillman, J.E., 1988. Mars global atmospheric oscillations: Annually synchronized, transient normal-mode oscillations and the triggering of
global dust storms. J. Geophys. Res. 93, 9433–9451.
Vanek, M.D., Wells, J.S., Maki, A.G., Burkholder, J.B., 1990. Heterodyne
frequency measurements on SO2 near 41 THz (1370 cm−1 ). J. Mol.
Spectrosc. 141, 346–347.
Wong, A.S., Atreya, S.K., Encrenaz, T., 2003. Chemical marker of possible hot spots on Mars. J. Geophys. Res. 108 (E4), 5026, doi:10.
1029/2002JE002003.
Wyatt, M.B., McSween Jr., H.Y., Chritensen, P.R., Head III, J.W., 2003.
Basalt, altered basalt, and andesite on the martian surface: Observations,
interpretations, and outstanding questions. In: Sixth International Conference on Mars, Pasadena. Abstract #3271 [CD-ROM].
Yung, Y.L., DeMore, W.B., 1999. Photochemistry of Planetary Atmospheres. Oxford Univ. Press, New York/Oxford.