CSHELL: Mapping of O2 dayglow, CO, and search for CH4

Icarus 190 (2007) 93–102 www.elsevier.com/locate/icarus

Long-term spectroscopic observations of Mars using IRTF/CSHELL: Mapping of O2 dayglow, CO, and search for CH4 Vladimir A. Krasnopolsky ∗,1 Department of Physics, Catholic University of America, Washington, DC 20064, USA Received 20 October 2006; revised 16 February 2007 Available online 12 March 2007

Abstract Long-term spectroscopic observations of the O2 dayglow at 1.27 µm result in a map of the latitudinal and seasonal behavior of the dayglow intensity for the full martian year. The O2 dayglow is a sensitive tracer of Mars’ photochemistry, and this map reflects variations of Mars’ photochemistry at low and middle latitudes. It may be used to test photochemical models. Long-term observations of the CO mixing ratio have been also combined into the seasonal–latitudinal map. Seasonal and latitudinal variations of the mixing ratios of CO and the other incondensable gases (N2 , Ar, O2 , and H2 ) discovered in our previous work are caused by condensation and sublimation of CO2 to and from the polar regions. They reflect dynamics of the atmosphere and polar processes. The observed map may be used to test global circulation models of the martian atmosphere. The observed global abundances of CO are in reasonable agreement with the predicted variations with the 11-year solar cycle. Despite the perfect observing conditions, methane has not been detected using the IRTF/CSHELL with a 3σ upper limit of 14 ppb. This upper limit does not rule out the value of 10 ppb observed using the Canada–France–Hawaii Telescope and the Mars Express Planetary Fourier Spectrometer. © 2007 Elsevier Inc. All rights reserved. Keywords: Mars, atmosphere; Spectroscopy; Atmospheres, compostion; Abundances, atmospheres

1. Introduction High-resolution spectroscopy of the planetary atmospheres from ground-based observatories is a traditional tool to study their chemical compositions. Using long-slit spectrographs, it is possible to map species distributions over a disk of a planet and study chemical and dynamical variations in its atmosphere. Here we report our long-term mapping observations of the martian atmosphere using the cryogenic echelle spectrograph CSHELL at the NASA Infrared Telescope Facility (IRTF) on Mauna Kea, Hawaii. Photochemical response of the martian atmosphere to changing insolation, dynamics, and thermal conditions on the planet * Corresponding address: 6100 Westchester Park Dr. #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, Science Mission Directorate, Planetary Astronomy Program. 0019-1035/$ – see front matter © 2007 Elsevier Inc. All rights reserved. doi:10.1016/j.icarus.2007.02.014

is among the basic problems in the studies of Mars. This problem may be solved by observations of the seasonal, latitudinal, and diurnal variations of photochemical tracers on Mars. The O2 dayglow at 1.27 µm, which is excited by photolysis of ozone and quenched by CO2 below ∼15 km on Mars, is one of these tracers. Our long-term measurements of the O2 dayglow are an important part of our observational program for the martian atmosphere and will be discussed in this paper. Study of variations of the CO mixing ratio on Mars is another part of our observational program. Our observations of CO on Mars were initially motivated by the reported variations of the CO mixing ratio with elevation along the slopes of the great martian volcanoes that questioned the basics of Mars’ photochemistry. Our observations mostly closed this problem and revealed a new phenomenon of seasonal variations of incondensable gases in the subpolar regions. This phenomenon related to condensation of CO2 and dynamics of the polar processes is currently the basic goal of our observations of CO on Mars. The measured seasonal and latitudinal variations of the CO mixing ratio as well as variations of the global CO abundance will be discussed below.

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The presence of methane in the martian atmosphere is the key problem that may be relevant to a possible life on Mars (Krasnopolsky, 2006a). We will report in this paper our attempt to detect methane on Mars and study its latitudinal variations using IRTF/CSHELL. This instrumentation is currently among the best to search for minor species by means of high-resolution spectroscopy. We believe that our results will be helpful for further studies of the problem. 2. Basic properties of IRTF/CSHELL CSHELL (Greene et al., 1993) is a long-slit spectrograph for the spectral range from 1.08 to 5.6 µm. Its detector is an InSb array of 256 spectral pixels by 150 spatial pixels. The array is cooled to 30 K. Each pixel is 0.2 × 0.2 arcsec2 and equals 9 × 10−6 ν0 in the dispersion direction. Here ν0 is a chosen central wavenumber which is determined by a combination of the grating and the circular variable filter positions. The spectral coverage is 256 × 9 × 10−6 ν0 = 0.0023ν0 . The instrument resolving power is ν/δν ≈ 4 × 104 for a slit width of 0.5 arcsec and corresponds to a resolution element of ∼3 pixels. The IRTF telescope diameter is 3 m. The CSHELL plus telescope point spread function has typically a full width at half maximum (FWHM) of ∼1 arcsec. The Mauna Kea summit has an elevation of 4.2 km. A low pressure of 0.6 bar and the dry atmosphere above the telescope with a typical water abundance of 2 pr. mm are very advantageous for infrared astronomy. An observing sequence with CSHELL involves observations of a target, atmospheric foreground (typically 30 arcsec off the target), flat field, dark current, a standard calibration star and a nearby star of 7–9th magnitude for focusing. The data reduction was described in Krasnopolsky and Bjoraker (2000) and Krasnopolsky (2003a, 2003b). 3. Mapping of the O2 dayglow at 1.27 µm Observations of the variations of Mars’ photochemistry with season, latitude, and local time are among the main objectives in the study of the martian atmosphere. This task requires longterm observations to cover the full martian year. We have found a sensitive tracer for this study and will discuss results of our observations in this section. The O2 (a 1 g → X3 g− ) dayglow at 1.27 µm was detected on Mars by Noxon et al. (1976) and later observed by Traub et al. (1979). The dayglow is excited by photolysis of ozone, and the O2 (1 ) molecules formed either emit photons at 1.27 µm or are quenched in collisions with CO2 . Clancy and Nair (1996) calculated the seasonal behavior of Mars’ photochemistry at low and middle latitudes. They argued that the atmosphere is dust free and therefore cold at aphelion and dusty and warm at perihelion. This changes significantly the condensation level for water vapor and strongly affects the water photolysis and production of odd hydrogen. Water and therefore odd hydrogen are strongly depleted above the condensation level, and this gives rise to ozone. The calculated ozone column abundances vary by a factor of 3 at low latitudes. Even

Fig. 1. IRTF/CSHELL spectrum of the O2 dayglow at 1.27 µm. The spectrum consists of 256 pixels of 0.072 cm−1 each. The main spectral features are the telluric O2 absorption lines and solar lines (S). The martian emission lines O2e are Doppler-shifted relative to the telluric lines. The solar lines are also Doppler-shifted. These lines are Fe 7892.30 cm−1 and Fe 7903.77 cm−1 . The line at 7897.35 cm−1 is present but not identified in Livingstone and Wallace (1991).

higher, by a factor of ≈10, are the variations of ozone above 20 km. Krasnopolsky (1997) found that the best way to observe high-altitude ozone is by measuring the O2 dayglow at 1.27 µm (Fig. 1) using CSHELL. This dayglow is quenched by CO2 below 15–20 km and therefore reflects high-altitude ozone, which is the most sensitive tracer for Mars’ photochemistry at low and middle latitudes. Krasnopolsky (1997) argued that the IRTF/CSHELL mapping of the O2 dayglow would provide a photochemical support to the MGS/TES observations of temperature profiles, H2 O, dust and ice aerosol opacities (Smith, 2002, 2004). Results of mapping of the O2 dayglow were discussed by Krasnopolsky and Bjoraker (2000) for LS = 68◦ and 112◦ , by Novak et al. (2002) for LS = 68◦ , and by Krasnopolsky (2003b) for 68◦ , 112◦ , 148◦ , and 173◦ . As usually, the areocentric longitude of the Sun LS specifies seasons on Mars with LS = 0◦ and 180◦ at northern spring and fall equinoxes, 71◦ and 251◦ at aphelion and perihelion, respectively. The IRTF/CSHELL observations of the O2 dayglow require a good combination of Mars angular diameter D  8 and geocentric velocity |V |  10 km/s, because the instrument resolution corresponds to the Doppler velocity of 7.5 km/s. These conditions are met for short periods before and after Mars oppositions. Some observations are lost because of bad weather and dust storms on Mars. Extraction of the weak O2 dayglow emission lines against the strong telluric O2 absorption lines and the intense continuum of the solar light reflected by Mars require some sophisticated techniques. Corrections for dark current, flat field, sky background, bad pixels, some misalignment between the instrument grating and detector, a possible extinction by cirrus, airmass, and ground reflection were discussed in Krasnopolsky

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Table 1 Observing conditions LS Date Heliocentric distance (AU) Geocentric distance (AU) Mars diameter (arcsec) Geocentric velocity (km s−1 ) Phase angle East longitude of CM Central latitude Local time at CM Mean local time

10◦ 10/02/06 1.581 1.154 8.1 17.3 38.5 261 −13.4 9:42 9:42

68◦ 21/01/97 1.666 0.968 9.7 −15.1 31.7 12 23.3 14:18 14:18

112◦ 20/03/99 1.624 0.746 12.6 −12.9 24.6 108 15.2 13:38 13:30

148◦ 20/04/01 1.541 0.734 12.8 −13.8 32.7 193 −1.4 13:59 12:30

173◦ 23/04/03 1.481 1.051 8.9 −14.4 42.7 30 −15.1 14:36 13:00

247◦ 10/07/05 1.382 0.949 9.9 −10.1 47.4 306 −21.5 15:25 15:25

312◦ 08/12/03 1.446 0.908 10.3 14.5 42.2 156 −26.4 9:00 11:30

Central meridian (CM) is given for the O2 dayglow observations. The other parameters do not significantly change during the observations. Mean local time is given for the latitudinal dependences of the O2 dayglow intensities in Fig. 2.

and Bjoraker (2000). Two other corrections, for the nonGaussian shape of the telluric O2 lines and for the instrument point spread function, were developed in Krasnopolsky (2003b) and used for the analysis of two new and reanalysis of two previous observations. The correction for the instrument point spread function is especially important at the limb and terminator where this correction is of a factor of 2. The retrieved and corrected vertical dayglow intensities and their variations with latitude and local time at four seasonal points were discussed in Krasnopolsky (2003b). Now we add to our database the O2 dayglow observations at three other seasonal points at LS = 247◦ , 312◦ , and 10◦ . With these points our observations cover rather uniformly the full martian year. The observing conditions are given in Table 1. The best spectral interval for the observations of the O2 dayglow with CSHELL is shown in Fig. 1. The O2 emission lines are typically weaker than those in the figure, and only four of the seven lines are accessible for the dayglow extraction at negative Doppler shifts. These lines are R1R1, R3R3, R3Q4, and R5Q6. The R5R5 and R7R7 lines add to this list at positive Doppler shifts. The other lines are either weak or contaminated by the solar lines. Latitudinal variations of the O2 dayglow intensity at seven seasons on Mars are shown in Fig. 2. The vertical intensity varies from 0.4 MR at 40◦ N for LS = 247◦ to 17 MR at 65◦ S for LS = 173◦ . (One MegaRayleigh is the emission of 1012 photons per cm2 s in 4π steradians.) The data in Fig. 2 refer either to the central meridians or to the local times between the noon and those for the central meridians. Local times at the central meridians and the mean local times for the curves in Fig. 2 are given in Table 1. The dayglow is maximum at 12:30, and the intensity variations from 9:00 to 15:00 are typically ∼20%. The martian atmosphere is optically thin at the wavelengths responsible for photodissociation of ozone. Therefore, the diurnal variations of the O2 dayglow reflect diurnal variations of the ozone abundances and vertical profiles while the photodissociation rate coefficient remains constant. Our observations of the diurnal variations of the O2 dayglow are considered in Krasnopolsky (2003b). The O2 dayglow intensities from Fig. 2 may be combined to build up a map of the dayglow variations with latitude and season (Fig. 3). This map reflects a photochemical response of

Fig. 2. Latitudinal variations of the O2 dayglow at 1.27 µm observed at seven seasons. The dayglow intensities at LS = 247◦ and 312◦ are scaled by a factor of 5.

the martian atmosphere to latitudinal and seasonal variations of temperature profiles, water vapor abundances, aerosol opacities, and dynamical processes in the atmosphere. This map may be used for predictions of the dayglow intensity at various observing conditions. It is also a test for modeling of Mars’ photochemistry and its variations. Both seasonal and interannual variabilities of thermal and dynamic properties of the martian atmosphere have been studied by MGS/TES (Smith, 2004). The interannual variability is typically low except for the southern spring LS = 180–270◦ . Global dust storms appear in this perihelion period and dramatically change the thermal balance in the lower atmosphere. Our seasonal–latitudinal map in Fig. 3 is applicable in this period to the conditions without global dust storms. The only published photochemical general circulation model (Lefevre et al., 2004) does not consider the O2 dayglow intensities. The recent model of variations of Mars’ photochemistry (Krasnopolsky, 2006b) is based on the MGS/TES data on T , H2 O, and aerosol abundances. Along with the standard gasphase chemistry, that model involves a weak loss of peroxide on the water ice aerosol particles. The model (Fig. 3) is generally in very good agreement with the observational data on the O2 dayglow and the ozone and peroxide abundances.

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Fig. 3. Seasonal–latitudinal variations of the O2 dayglow at 1.27 µm (in MR). Observations are in the lower panel. The upper panel shows the results of photochemical modeling (Krasnopolsky, 2006b) in the latitude range covered by the MGS/TES observations of H2 O (Smith, 2004).

The model was calculated for the mean daytime conditions at the latitudes covered by the MGS/TES observations of H2 O. It extends to 50◦ S at LS = 173◦ and cannot show a very strong emission centered at 65◦ S in this season. However, the northern wing of this emission extends to 35◦ S, and the observed intensity near 40◦ S exceeds the calculated value. Maybe the observed strong emission is an occasional phenomenon that was caused by specific weather conditions. It is necessary to continue the O2 dayglow observations to study interannual variability and get a more detailed map. The O2 dayglow at 1.27 µm is now observed by the SPICAM IR spectrometer at the Mars Express orbiter, and the data for the first martian year were recently published (Fedorova et al., 2006). Generally, the SPICAM observations provide the regular monitoring of the O2 dayglow and better coverage of high latitudes that are typically poorly seen by groundbased telescopes. The IRTF/CSHELL observations are made with a mean rate of once per year and the spatial resolution which is significantly lower than that for the spacecraft data. Obviously the SPICAM observations are not affected by the telluric absorption that restricts the ground-based observations to periods with high geocentric velocity. Advantages of

IRTF/CSHELL are the higher spectral resolution (by a factor of 18) and the high signal-to-noise ratio (∼300 for the observations in 2005–2006). The IRTF/CSHELL observations may be better for the weak O2 dayglow typical of the low latitudes and especially near Mars’ perihelion. These observations are not linked to a certain local time as in the case of the Mars Express data. Fedorova et al. (2006) made a detailed comparison of the SPICAM results with the observations in Krasnopolsky (2003b). The agreement is generally reasonable. Significant differences at LS = 68◦ are properly explained by the very different local time (7 and 14.3 h for the SPICAM and IRTF observations, respectively) and those for LS = 148◦ at 35–50◦ S by interannual variations. The published SPICAM observations cover one martian year and cannot show interannual variability as well as our data. Our new data are also generally consistent with the SPICAM observations made at the similar conditions. Overall the interannual variations in the O2 dayglow from the comparison of the two datasets appear to be higher than those observed by MGS/TES for T , H2 O, and aerosol opacities out of the global dust storms.

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4. Mapping of the CO mixing ratio on Mars Mapping of the CO mixing ratio on Mars is another part of our observational program for Mars. Our initial goals were variations of CO with local time and elevation to test the predictions of photochemistry and some controversial observational data. Later our observations revealed a new phenomenon of seasonal variations of the CO mixing ratio in the subpolar regions, and observations of these variations are currently the main objective of our study. CO is the basic product of the CO2 photolysis. It reacts very slowly with O, O2 , and O3 ; therefore the CO mixing ratio is expected at 0.08 in the dry CO2 atmosphere (Nair et al., 1994). The observed values are smaller by two orders of magnitude. This so-called problem of the stability of the martian CO2 atmosphere was generally solved 35 years ago using a very effective catalysis by products of the H2 O photolysis. Furthermore, this catalysis is so effective that the predicted CO abundances are smaller than the mean observed value, 8 × 10−4 , by a factor of ∼5. To remove this disagreement, Nair et al. (1994) and Krasnopolsky (1995) suggested some changes in the standard gas-phase chemistry that, however, have not been confirmed by the later laboratory studies. Recently Krasnopolsky (2006b) proved that the situation cannot be solved using heterogeneous loss of odd hydrogen on the water ice aerosol, and the problem remains unsolved. The CO production rate is just equal to the flux of the solar photons with λ < 200 nm that dissociate CO2 . This global and annually mean flux is 9 × 1011 cm−2 s−1 , the CO column abundance is 1.8 × 1020 cm−2 for the mixing ratio of 8 × 10−4 , and a ratio of these values is the CO mean lifetime, 6 years. This lifetime is much longer than the martian day; therefore no variations of CO with local time are expected. Time of vertical mixing is H 2 /K ≈ 2 weeks in the lower atmosphere; therefore the CO mixing ratio should be constant up to ∼40 km. Here H is the scale height and K ≈ 106 cm2 s−1 (Korablev et al., 1993) is the eddy diffusion. Our observations were initially motivated by puzzling results from the low resolution spectroscopy (λ/δλ ≈ 70) of the CO (2–0) band at 2.35 µm from the Phobos orbiter (Rosenqvist et al., 1992). The retrieved CO mixing ratios over the great martian volcanoes at elevation of ∼20 km were smaller than that above planitia by a factor of 5–8, that is, completely inconsistent with the predictions of photochemistry. Our observations on March 20, 1999 (Krasnopolsky, 2003a) did not reveal any variation of the CO mixing ratio with surface elevation. Evidently we could not resolve the martian volcanoes using IRTF/CSHELL; however, no significant variation was observed in the range from −6 to 3 km. Hunten (1993) pointed out that small changes in the baseline for the CO band could remove the variation of the CO mixing ratio in the Phobos observations. No variation of CO with local time was observed by Krasnopolsky (2003a) as well. A new phenomenon revealed in Krasnopolsky (2003a) was the observed increase in the CO mixing ratio by a factor of ∼1.5 from the equator to 50◦ S. The observations were made during the southern winter at LS = 112◦ , and this increase was prop-

Fig. 4. IRTF/CSHELL spectrum of Mars near 1.57 µm. The spectrum consists of the CO and CO2 lines and the solar lines (S). Asterisks mark the lines used for the extraction of the CO mixing ratio. The wavenumber scale is corrected for the Doppler shift, and the whole spectrum including the solar lines is at rest position.

erly explained by condensation of CO2 at the winter polar cap. Condensation and sublimation of CO2 result in an enrichment and depletion of incondensable long-living gases (N2 , Ar, O2 , CO, H2 ) on the polar caps and nearby regions. Krasnopolsky (2003a) argued that the value and the extent of this effect are directly related to dynamics of the atmosphere and polar processes and should be simulated by the general circulation models (GCM). This effect was predicted by Krasnopolsky (1993) and was later observed for Ar using the gamma ray spectrometer on the Mars Odyssey orbiter (Sprague et al., 2004). Our further mapping observations of CO are aimed at the seasonal variations of this effect. A typical CSHELL spectrum of Mars near 1.57 µm is shown in Fig. 4. The spectrum includes a few weak lines of CO and CO2 that make it possible to derive the CO mixing ratio. Some sources of error significantly cancel out in this mixing ratio. We also tried to use other spectral intervals within the CO (3–0) band. However, these attempts were less successful. For example, we found that the P6 line at 6325.80 cm−1 , which was used by Kaplan et al. (1969) for the first determination of the CO mixing ratio on Mars, is contaminated by the CO2 line at 6325.77 cm−1 . The contribution of the CO2 line to the total absorption varies from 10% at 210 K to 50% at 270 K. Therefore we choose only the observations in the interval of 6379–6394 cm−1 for our analysis. The data processing for extraction of the CO mixing ratio from the observed spectra was described in detail in Krasnopolsky (2003a). That technique was based on an effective temperature which was derived from the MGS/TES temperature profiles (Smith, 2004) for each line depending of its strength. Now we use the MGS/TES temperature profiles directly to sum up the absorption effect in the atmosphere divided by 30 layers. Fig. 5 shows the CO mixing ratio as a function of surface elevation from the observations on December 8, 2003. The range of elevations is from −4.5 to 8.5 km, and no variation is seen.

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Fig. 5. Mars’ CO mixing ratio (corrected for the latitudinal dependence) versus surface elevation for the observations on December 8, 2003. The dashed line is a linear fit to the data. The statistical uncertainties are low because of a large number (728) of the observed spectra.

Variations of the CO mixing ratio (fCO ) with latitude at four seasons are shown in Fig. 6. We will analyze this figure using the simulations of the polar cap dynamics by Pollack et al. (1993) and Hourdin et al. (1995). The simplest case is for LS = 112◦ with fCO constant in the northern hemisphere and increasing from the equator to the south polar cap. The CO2 condensation rate is maximal near LS ≈ 100◦ , and the CO asymmetry delays for some time being near its maximum at LS = 112◦ . A similar situation but for the northern winter is at LS = 312◦ . However, the northern winter is near Mars’ perihelion, it is not so cold as the southern winter, and the CO2 condensation rate and the CO enhancement are significantly weaker than that on the south at LS = 112◦ . Sublimation of the south cap ended a month before the observations, and the decrease in fCO to the south has not been completely removed by mixing. The south cap is starting to sublime at LS = 173◦ . The meridional mixing has made fCO flat up to 45◦ S for a month after the end of condensation. The other features are similar to those at LS = 112◦ . Explanation of the CO behavior at LS = 10◦ requires some complicated dynamics.

The CO distributions for four seasons are generally insufficient to build a map of the latitudinal and seasonal variations, and we consider the map in Fig. 7 as a preliminary version. Fortunately, the observations are spaced rather uniformly over the martian year. We hope to improve this map by adding future observations. Observations of Hellas using the low resolution (λ/δλ ≈ 100) imaging spectrometer Omega at the Mars Express orbiter revealed seasonal variations of the CO mixing ratio from (8 ± 4) × 10−4 at LS = 340◦ to (28 ± 8) × 10−4 at LS = 132◦ (Encrenaz et al., 2006). Krasnopolsky (2003a) studied variation of the CO mixing ratio over the slope of Hellas with a vertical drop of 9 km. The observed variation was small, and fCO at Hellas was similar to that in the neighbor regions. Hellas is at 43◦ S, and our observations at this latitude give fCO = 8, 15, 15.5, and 9 × 10−4 at LS = 10, 112, 173, and 312◦ , respectively. Our minimum value is similar to that in Encrenaz et al. (2006). Our maximum which is expected at (18 ± 5) × 10−4 is smaller but the uncertainties are overlapping. 5. Solar cycle variations of the CO abundance The problem of variations of the CO abundance with the 11year solar cycle was considered by Krasnopolsky (1993). The column CO2 photolysis rate varies by 25% within the solar cycle. The abundance of water vapor and its photolysis rate are insensitive to solar activity. There are some fine mechanisms that would increase the variations of the CO abundance up to a factor of ∼4. However, the mean lifetime of CO in the martian atmosphere is 6 years and comparable with the period of the solar cycle. Therefore the expected variations of CO with the solar cycle are equal to 35% with a phase shift of 2 years. Our observations show that the CO mixing ratio varies with season and latitude, and the observed variations are comparable with the expected solar-cycle variations. Data from different ob-

Fig. 6. Latitudinal variations of the CO mixing ratio at four seasons. The observational data are approximated by broken lines.

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Fig. 7. Seasonal–latitudinal variations of the CO mixing ratio (in 10−4 ).

servations of CO since its first detection by Kaplan et al. (1969) are not helpful because of the significant uncertainties in the observations (∼25%) and possible systematic errors. Our observational data are currently the best to study the problem. They cover a rather long period of seven years, involve latitudinal distributions, and have been uniformly processed with similar possible systematic errors for each period. Using the MGS/TES temperature profiles (Smith, 2004), the MOLA elevations, the Viking measurements of the seasonal variations of pressure (Hess et al., 1980), and the global and seasonally mean pressure of 6.1 mbar, we may calculate distributions of the surface pressure p over the globe for each of the four seasons (Krasnopolsky, 2006b). Then the global abundance of CO is A CO = 2πR μg

π/2

2

p(θ)fCO (θ ) cos θ dθ. −π/2

Here R = 3.39 × 108 cm is Mars’ radius, A = 6.022 × 1023 is the Avogadro number, μ = 43.5 is the mean molecular mass, and g = 371 cm s−2 is the gravity acceleration. To calculate the global CO, our observations and the MGS/TES data are extrapolated to the poles. A polar cap θ > 70◦ is equal to 3% of the globe surface, and an error associated with this extrapolation is not large. The calculated global CO abundances at various times of the observations are shown in Fig. 8. The best fit to the observed values by a function   2π CO = a sin (t − t0 − b) + c T is shown in Fig. 8. Here T = 11 years is the period, t0 = 1996 is the time of the last solar minimum, and a, b, and c are the fitting parameters. This fitting results in the CO variations by a factor of (c + a)/(c − a) = 1.21 with a time delay of 1.24 years. The uncertainties of the global CO abundances are comparable

Fig. 8. Long-term variations of the total CO abundance in the martian atmosphere. The observed abundances are fitted by a sinusoid with the period of the solar cycle (11 years). The last solar minimum was in 1996.

to the observed CO variations, the number of the observations is comparatively low, and the CO photochemistry is poorly reproduced by the models. Taking into account all these facts, the agreement between the observations and the calculations by Krasnopolsky (1993) may be considered as reasonable. 6. Search for methane We have attempted to detect methane on Mars using the CSHELL spectrograph at NASA IRTF on February 10, 2006. Conditions of the observation were the best in terms of (1) Mars’ geocentric velocity and its sign, (2) position of Mars on the sky near the zenith, (3) very low humidity, and (4) very good seeing. We observed the R0 and R1 lines of the strongest band of CH4 at 3.3 µm. Their wavenumbers are 3028.75 and 3038.50 cm−1 , respectively.

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Mars’ geocentric velocity was at its maximum value of 17.3 km s−1 and provided the highest Doppler shift. The negative Doppler shift is better for detection of methane because the left wings of the R0 and R1 telluric lines are rather clean at the positions of the martian lines while the right wing of R0 is contaminated by the H2 O line at 3028.91 cm−1 and the CH4 line at 3028.85 cm−1 . The right wing of R1 is contaminated by the CH4 lines at 3038.61 and 3038.63 cm−1 . Telluric absorption in the H2 O line is equal to 0.04 for the mean humidity of 2 pr. mm above Mauna Kea, and the absorptions from each of the three CH4 lines are ∼0.25. The telluric airmass factor varied during the observation of CH4 from 1.01 to 1.15, minimizing the effect of the telluric absorption. Humidity was 0.5 pr. mm above Mauna Kea during the observation, far below the mean value. We made three observing sequences for the R0 line and two sequences for the R1 line. Each sequence involved 8 oneminute exposures of Mars and flats, darks, and foreground observations 30 arcsec off Mars. We summed up the processed observations in each sequence and got three spectral frames for R0 and two frames for R1. The instrument slit was placed at the central meridian during the observations. Mars angular diameter was equal to 8.1 arcsec, the pixel size is 0.2 × 0.2 arcsec2 , and each frame includes 40 spectra of Mars. A spectrum is a row of 256 pixels of ∼0.03 cm−1 each. Wavenumbers for each pixel in a column of the detector array are not the same but vary within ∼0.05 cm−1 . This may be a significant source of error if all 40 spectra are summed up to reduce the noise. To eliminate this error, we transform the data from each pixel to eight sampling points. The simplest way to do this is the linear interpolation between the neighbor pixels. More complicated cases apply the formulae of Newton, Stirling, and Bessel. However, all these methods fix the central sampling point at the pixel reading while fixing the sum of eight sampling points at the pixel reading is more appropriate to detectors with small gaps between pixels. We have developed a method to apply this approach using the parabolic interpolation. Finally we get five frames of 2000 × 150 points with a fixed wavenumber in each column. One of the observed spectra centered at the R0 line is shown in Fig. 9. The spectrum consists of telluric absorption lines of H2 O, CH4 , and O3 . A spectrum of the foreground 5 arcsec off the limb of Mars from the same frame is also shown, and the telluric lines are seen in emission in this spectrum. Absorption of monochromatic light I from an external source in the Earth’s atmosphere is accompanied by thermal emission: dI = −I (τ ) dτ + B(τ ) dτ. Here τ is the slant optical depth in the Earth’s atmosphere and B(τ ) is the blackbody thermal emission at temperature T (τ ). Integration of this differential equation results in I (τ ) = I0 e

−τ

+e

−τ

τ B(t)et dt. 0

The second term is the foreground emission, and subtraction of the foreground spectrum from the target spectrum corrects the latter for the thermal emission in the Earth’s atmosphere.

Fig. 9. One of the spectra centered at the R0 line and observed near the equator. Positions of telluric H2 O and CH4 lines with strengths exceeding 10−24 and 10−22 cm, respectively, are shown. The foreground spectrum observed 5 arcsec off the Mars limb is shown in the lower part of the plot.

Fig. 10. The spectrum from Fig. 9 corrected for the foreground and fitted by a synthetic spectrum (dotted line). Longitudes are 62◦ –72◦ W. The difference between the observed and synthetic spectra is ∼0.5%; therefore the synthetic spectrum is poorly seen except the positions of the martian H2 O lines. These differences for sums of five spectra centered at the given latitudes are shown scaled by a factor of 3. The mean difference and the bar of ±3σ are shown by solid lines. Thin vertical lines correspond to the position of the martian CH4 line at FWHM. A 3σ upper limit to CH4 is equal to 22 ppb for the mean curve.

We have found that subtraction of the interpolated foreground spectrum from the same frame is more accurate than using the foreground frames measured 30 arcsec off the basic frame. The corrected spectra are grouped in seven bands of five spectra each for further reduction of the noise. This gives seven averaged spectra displaced by one arcsec in each frame. The mean R0 spectrum observed near the equator is shown in Fig. 10. This spectrum is fitted by a synthetic spectrum with variable abundances of H2 O, CH4 , and O3 in the terrestrial atmosphere. The line parameters for fitting are taken from the HITRAN 2004 spectroscopic database. We use temperature and pressure profiles in the Earth’s atmosphere from the US Standard Atmosphere 1976. Using the pure collisional and Voigt line profiles give essentially the same results. The synthetic

O2 dayglow, CO, and CH4 on Mars

Fig. 11. The ATMOS spectrum of the Sun (Farmer and Norton, 1989) convolved with the CSHELL resolution.

spectra are finally convolved using a Gaussian with a width which is also a fitting parameter. The best solar spectrum in the range of the R0 and R1 lines was observed by the ATMOS orbiter (Farmer and Norton, 1989) with resolution of 0.01 cm−1 . The spectrum looks flat and uniform in this atlas near the R0 and R1 lines because all features in the spectrum are of a few percent. However, these small features may significantly affect our study, and we apply the ATMOS spectrum from its file. This spectrum convolved by the CSHELL resolution is shown in Fig. 11. The solar spectrum reflected by Mars is affected by (1) the heliocentric velocity of Mars that was equal to 2.0 km s−1 , (2) a projection of Mars rotation velocity on the direction to the Sun, and (3) Mars velocity relative to the observer of 17.3 km s−1 . The second component varies from the equator to the poles with a mean value of −0.1 km s−1 . The total Doppler shift for the solar spectrum is −0.193 cm−1 . The coincidence between the observed and synthetic spectra is so good that the synthetic spectrum is poorly seen in Fig. 10. The difference between these spectra demonstrates four Doppler-shifted lines of the martian water vapor. The observed Doppler shift is −0.173 ± 0.003 cm−1 and agrees with the geocentric velocity of 17.3 km s−1 . The line widths confirm the CSHELL resolving power ν/δν ≈ 4 × 104 . There are no absorptions of the martian methane in the difference spectra at various latitudes as well as in the mean difference spectrum. The mean difference is 0.00813 in the interval of ±0.4 cm−1 centered at the R0 line in the mean difference spectrum. Transmittance at the position of the martian line is 0.5; therefore σ = 0.0163. This σ refers to a true pixel size p = 0.0294 cm−1 . Resolution element δν is a full width at half maximum of the instrument response function which is adopted as a Gaussian δν −1 0 2 exp(− 12 ( ν−ν d ) ); then d = 2(2 ln 2)1/2 = 0.0322 cm . The total equivalent width of a line is w = I0 d(2π)1/2 = I0 × 2.75p. Noise is independent in the neighbor pixels; therefore a 3σ upper limit to the line equivalent width is w3σ = 3 × 0.0163 × 2.751/2 × 0.0294 cm−1 = 2.38 × 10−3 cm−1 . The R0 line strength is 1.65 × 10−19 cm at 200 K and the mean two-way

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Fig. 12. Same as Fig. 10 for the R1 line. Longitudes are 72◦ –82◦ W. A 3σ upper limit to CH4 is equal to 17 ppb for the mean curve. The retrieved abundances of the telluric H2 O, CH4 , and O3 agree with other data on these species in the terrestrial atmosphere. Differences between the abundances in Fig. 10 and this figure reflect uncertainties of the abundances and their variations for 40 min between the observations.

martian airmass along the slit is 2.81; therefore a 3σ upper limit to the vertical CH4 abundance on Mars is 5.14 × 1015 cm−2 . The mean atmospheric pressure is 6.2 mbar calculated using the Viking 1 measurements (Hess et al., 1980) and the global and annual mean pressure of 6.1 mbar. Then the column density of the atmosphere is 2.31 × 1023 cm−2 and the 3σ upper limit to the methane mixing ratio on Mars is 22 ppb. A similar calculation for the R1 line (Fig. 12) results in the upper limit of 17 ppb. Weighted summing the mean differences at ν0 ± 0.4 cm−1 for the R0 and R1 lines gives a final 3σ upper limit to the CH4 mixing ratio on Mars of 14 ppb. Inspection of the difference spectra for various latitude in Figs. 10 and 12 shows significant similarities between the spectra. These similarities indicate some systematic errors that cannot be reduced by coadding. These systematic errors dominate in the mean difference spectra, and further coadding of the frames does not improve the uncertainty. Therefore we do not see any means to improve our upper limit using the same instrumentation. Generally this upper limit does not rule out the value of 10 ppb observed by Krasnopolsky et al. (2004a, 2004b) and the PFS team (Formisano et al., 2004). 7. Conclusions Our long-term spectroscopic observations of the O2 dayglow at 1.27 µm result in a map of the latitudinal and seasonal behavior of the dayglow intensity. The O2 dayglow is a sensitive tracer of Mars’ photochemistry, and this map reflects variations of Mars’ photochemistry at low and middle latitudes. It may be used to test photochemical models. Long-term observations of the CO mixing ratio have been also combined into the seasonal–latitudinal map. Seasonal and latitudinal variations of the mixing ratios of CO and the other incondensable gases (N2 , Ar, O2 , and H2 ) are caused by condensation and sublimation of CO2 to and from the polar regions.

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They reflect dynamics of the atmosphere and polar processes. The observed map may be used to test global circulation models of the martian atmosphere. Fitting of the observed global abundances of CO by a sinusoid with the solar cycle period of 11 years results in the solar cycle variations of CO by a factor of 1.2 with a time delay of 1.2 years. Though these values are rather uncertain, they are in reasonable agreement with the model predictions. Despite the perfect observing conditions, methane has not been detected using the IRTF/CSHELL with a 3σ upper limit of 14 ppb. This upper limit does not rule out the value of 10 ppb observed using the Canada–France–Hawaii Telescope (Krasnopolsky et al., 2004a, 2004b) and the Mars Express Planetary Fourier Spectrometer (Formisano et al., 2004). We do not see any means to improve this limit using IRTF/CSHELL. Acknowledgments This work is supported by the NASA Planetary Astronomy Program (Grant NNG04GI32G to V.A. Krasnopolsky). The author acknowledges the support and help from the IRTF staff and its director Alan Tokunaga. References Clancy, R.T., Nair, H., 1996. Annual (perihelion–aphelion) cycles in the photochemical behavior of the global Mars atmosphere. J. Geophys. Res. 101, 12785–12790. Encrenaz, T., Fouchet, T., Melchiorri, R., Drossart, P., Gondet, B., Langevin, Y., Bibring, J., Forget, F., The OMEGA team, 2006. Seasonal variations of the martian CO over Hellas as observed by Omega/Mars Express. Bull. Am. Astron. Soc. 38, 600. Farmer, C.B., Norton, R.H., 1989. A high-resolution atlas of the infrared spectrum of the Sun and the Earth atmosphere from space. I. The Sun. NASA Ref. Publication 1224, vol. 1. Fedorova, A., Korablev, O., Perrier, S., Bertaux, J.L., Lefevre, F., Rodin, A., 2006. Observations of O2 1.27 µm dayglow by SPICAM IR: Seasonal distribution for the first martian year of Mars Express. J. Geophys. Res. 111, doi:10.1029/2006JE002694. E09S07. Formisano, V., Atreya, S., Encrenaz, T., Ignatiev, N., Guiranna, M., 2004. Detection of methane in the atmosphere of Mars. Science 306, 1758–1761. Greene, T.P., Tokunaga, A.T., Toomey, D.W., Carr, J.S., 1993. CSHELL: A high spectral resolution echelle spectrograph for the IRTF. In: Proc. SPIE, vol. 1946, pp. 313–323. Hess, S.L., Ryan, J.A., Tillman, J.E., Henry, R.M., Leovy, C.B., 1980. The annual cycle of pressure on Mars measured by Viking Landers 1 and 2. Geophys. Res. Lett. 7, 197–200. Hourdin, F., Forget, F., Talagrand, O., 1995. The sensitivity of the martian surface pressure and atmospheric mass budget to various parameters: A comparison between numerical simulations and Viking observations. J. Geophys. Res. 100, 5501–5523. Hunten, D.M., 1993. CO on Mars: Comment on paper by Rosenqvist et al. Icarus 101, 42–44.

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