Space Science Reviews

, Volume 182, Issue 1–4, pp 19–84 | Cite as

Lower and Upper Ionosphere of Mars



The ionosphere of Mars has been explored mostly with the radio occultation experiment onboard Mariners 6, 7, 9; Mars 2, 3, 4, 6; Viking 1, 2, and more recently on Mars Global Surveyor (MGS) and Mars Express (MEX). In addition to the radio occultation experiment, MEX also carried Mars Advanced Radar for the Subsurface and Ionosphere Sounding (MARSIS) experiment which provided electron density profiles well above the main ionospheric peak. The atmosphere of Mars was measured directly by the neutral mass spectrometer onboard Viking 1 and 2 Landers. Later, an accelerometer and radio occultation experiment on MGS provided large data sets of atmospheric density at various locations in the upper and lower atmospheres of Mars, respectively. In this paper we review results of these upper and lower atmospheric/ionospheric measurements. Results of these measurements have been compared with theoretical models by several workers; therefore, we also review various atmospheric and ionospheric models of Mars.


Ionosphere of Mars 

1 Introduction

Broadly, the Martian atmosphere can be divided into two regions, namely (1) the lower atmosphere in which turbulent mixing is dominant so that the mean molecular mass of atmospheric constituents is almost independent of height and (2) the upper atmosphere in which molecular diffusion dominates and the constituents are separated according to their molecular mass. The atmosphere also contains the ionosphere, which is the result of ionization of neutral atmosphere by photons and particle impact. Mars’ upper atmosphere has been observed by Mariners 6, 7, 9; Mars 4, 5; Viking 1 and 2, Mars Global Surveyor (MGS) and Mars Express (MEX), but there have been very few measurements in the lower atmosphere of Mars. In this paper we review experimental results obtained on the lower and upper atmosphere of Mars from the various planetary missions. The understanding of complex behavior of Martian atmosphere and ionosphere requires a balanced effort in the area of theoretical modeling. Therefore we also review various atmospheric models of Mars, which have been developed during the last four decades.

Results from various planetary missions have been reviewed in the past by Whitten and Colin (1974), Schunk and Nagy (1980), Cravens and Nagy (1983), Mahajan and Kar (1988), Kar (1996), and Nagy et al. (2004). Recently Withers (2009) has reviewed variability in the dayside ionosphere of Mars while Haider et al. (2011) have reviewed experimental and modeling results in the ionosphere of Mars. But these reviews covered mainly Mars’ upper ionosphere. In this work we will review results on the upper and lower atmosphere and ionosphere as obtained from various missions as well as various theoretical models. Future missions like Mars Atmosphere and Volatile Evolution (MAVEN), Mangalyaan, ExoMars and Mars Exploration with Lander and Orbiters (MELOS) are in the pipeline to explore Mars environment. The objective of these missions is to obtain new science on the lower and upper atmosphere of Mars. In view of these great forthcoming missions, a review paper on the upper and lower atmosphere/ionosphere of Mars is desirable.

The present review consists of nine chapters. The first chapter gives a general introduction on the contents of this paper. In the second chapter we review Viking 1/2 and MGS measurements of density and temperature of the upper atmosphere of Mars. This chapter also describes exosphere, neutral chemistry and dynamics of Martian atmosphere. In the third chapter we review upper ionospheric measurements as well as various physical and chemical processes involved in the formation of Mars ionosphere. In particular, ionization sources like solar EUV, X-ray and particle radiations are discussed in this chapter. Observations on the upper ionosphere of Mars during disturbances like solar flares, solar energetic particles (SEP) and Coronal Mass Ejections (CMEs) are reviewed in chapter four. The modeling of the upper ionosphere during quiet and disturbed conditions is also reviewed in this chapter. In the fifth chapter we review measurements of density, temperature, pressure and dynamics of the lower atmosphere. The sixth chapter focuses on the lower ionospheric measurements accomplished by Mars 4/5, MGS and MEX. This chapter also includes composition and chemistry of meteoric ions. The seventh chapter reviews modeling results obtained on ion production rates, conductivity and ion concentrations due to impact of cosmic rays in the presence as well as in the absence of dust storms in the lower ionosphere of Mars. Knowledge of ion composition and chemistry of dust present in the Martian troposphere is required to understand the atmospheric electricity and wave propagation problem. In chapter eight, we discuss briefly those measurements, which at present are missing but are necessary to understand the complete science of the lower and upper atmosphere of Mars. Finally, in chapter nine we provide a summary and conclusions on the work reviewed in this paper. The objectives of future Mars missions, which will provide the missing information, are also highlighted in this chapter.

2 Upper Atmosphere of Mars

The early deep space missions like Mariners 4, 6, 7, 9 (Barth et al. 1971), Mars 4, 6 (Moroz 1976) and Viking 1, 2 (Nier and McElroy 1977) brought a dramatic improvement in our knowledge of the Mars atmosphere. These missions generated important data sets which helped to define the structure and properties of the Martian atmosphere. We have learned from these missions that the upper atmosphere of Mars consists of a thermosphere and an exosphere, and that the temperature in its upper atmosphere increases at a rate slower than that in the Earth’s upper atmosphere. This shows lower heating through the absorption of UV energy in the upper atmosphere of Mars than in the Earth’s atmosphere. Because of relatively small scale heights, the exobase for Mars occurs at a much lower altitude (200 km) than that for Earth (∼500 km). Mars has a much smaller gravity and is therefore subject to non-thermal escape processes. As a result, a large part of nitrogen and oxygen has escaped from its atmosphere (Hunten 1990).

2.1 The Earliest Measurements of Neutral Density and Temperature

The first direct measurements of the neutral atmosphere of Mars were accomplished by the two Viking Landers. Viking 1 landed on Mars on 20 July 1976 at 22.5N, 48W about 4 h after local noon with solar zenith angle of 44. Viking 2 landed on Mars at 48N, 22W on 3 September 1976 about 10 h after local midnight with a solar zenith angle of 44. The neutral mass spectrometer onboard Viking 1 and 2 measured the neutral composition during their landings. Figures 1A and 1B show the number densities of CO2, N2, CO, O2 and NO as measured by Viking 1 and 2 respectively (Nier and McElroy 1977). An important result derived from these measurements were the abundance of each of these gases relative to all other gases (known as mixing ratio) at the altitude below which turbulent mixing dominates (known as turbopause). The mixing ratios of these gases were CO2, 95.5 %; N2, 2.7 %; Ar, 1.5 %; CO, 0.4–1.4 %; O2, 0.17 %; and NO, 0.008 %. The values of the eddy mixing coefficients deduced at the turbopause were 5×107 cm−2 s−1 and 1.5×108 cm−2 s−1 for Viking 1 and 2, respectively. The turbopause altitude was located by these observations to be at 125 km. Nier and McElroy (1977), after fitting the best straight lines to the data above 140 km, found neutral temperatures of about 179, 175 and 164 K for CO2, N2 and O2, respectively for Viking 1 measurements. In the case of Viking 2 measurements corresponding to temperatures of these neutrals were 128, 111 and 110 K, respectively. The neutral temperature profiles of the Martian atmosphere above 120 km observed by Viking 1 and 2 are represented in Figs. 2A and 2B, respectively and are compared with those reported by Seiff and Kirk (1977) for altitudes below 120 km. The atomic oxygen density could not be measured by the mass spectrometer on Viking 1 and 2. However, it was inferred from \(\mathrm{O}_{2}^{+}\) and \(\mathrm{CO}_{2}^{+}\) densities measured by retarding potential analyser (RPA) experiment on Viking 1 and 2 (Hanson et al. 1977). Helium has also been detected on Mars by Krasnopolsky et al. (1994) through airglow measurements.
Fig. 1

(A) Neutral densities of CO2, N2, CO, O2 and NO observed from Viking 1; (B) neutral densities of CO2, N2, CO, O and NO observed from Viking 2 (from Nier and McElroy 1977)

Fig. 2

(A) Neutral temperature for Martian atmosphere above 120 km as measured by Viking 1. Temperatures obtained by Seiff and Kirk (1977) are plotted for comparison. (B) Same as A but for Viking 2. Uncertainties are also shown in both figures (Nier and McElroy 1977)

2.2 MGS Measurements of Atmospheric Waves and Density

Neutral mass spectrometers onboard Viking 1 and 2 Landers observed atmospheric compositions at two fixed locations. However, mass density in the upper atmosphere of Mars was measured by the accelerometer experiment onboard MGS during aerobraking period at different latitudes, longitudes, local solar times and seasons (Keating et al. 1998). Aerobraking took place in two phases—Phase 1 and 2. Phase 1 included orbits # P1-P201, from mid September 1997 to late March 1998, while Phase 2 included the orbits # P574-P1283 from mid September 1998 to early February 1999. Accelerometer data have been used extensively to study the dynamics of the upper atmosphere of Mars (e.g. Bougher et al. 2001; Withers et al. 2003; Krymskii et al. 2003; Haider et al. 2006, 2010a). Figure 3 shows the orbital distribution of mass densities at several altitudes (110, 130, 140, 150 and 160 km) at the low latitude region (0–25N) in Phase 2 of aerobraking of MGS mission during orbits P790 to P910 for the month of December 1998 (Haider et al. 2010a). At that time Mars had summer (Ls=90) with moderate solar activity period (f10.7=124). These densities were measured at solar zenith angle of 78 for inbound legs of 120 orbits. The data are not available at 110 km for orbits from P866 to P910. There is a large variation in densities in orbits P790 to P864 relative to that observed between orbits P865 and P910.
Fig. 3

MGS accelerometer measurements of mass densities for different orbits at different altitudes for low northern latitudes region (0–25N) (from Haider et al. 2010a)

2.3 Exosphere of Mars

Mars has a weak gravity, which allows an extended corona of hot species and escape of its lighter constituents in the exosphere. Although no measurements have been made in the exosphere of Mars, several theoretical 1-D models, over the past two decades, predicted hot oxygen corona population in the exosphere of Mars (e.g. Ip 1990; Lammer and Bauer 1991; Fox and Hac 1997; Kim et al. 1998; Lammer et al. 2000; Nagy et al. 2001; Fox and Yeager 2006; Cipriani et al. 2007; Fox 2009). The hot oxygen is produced in the exosphere of Mars due to dissociative recombination of \(\mathrm{O}_{2}^{+}\) ions. The escape flux of this gas can be estimated by counting the number of dissociative events in the exosphere. The determination of the current escape of the upper atmosphere/ionosphere allows us to understand the history of Mars atmosphere, climate, liquid water and planetary habitability.

The structure of the Martian upper atmosphere is complicated and therefore can’t be described globally by 1-D thermosphere/ionosphere models. Bougher et al. (2009) and Valeille et al. (2009a) have developed 3-D Mars Thermosphere Global Circulation Model (MTGCM) and 3-D Direct Simulation Monte Carlo (DSMC) models respectively to study the variation of thermosphere and exosphere of Mars at different local times, latitudes and longitudes. A combination of these models describes self consistently the Martian upper thermosphere/ionosphere and exosphere globally (Valeille et al. 2009b). Figure 4 represents density maps of O and CO2 in northern and southern hemisphere of Mars at an altitude of 190 km (Valeille et al. 2009b). Atomic oxygen reaches its maximum densities at low latitudes on the nightside near the morning terminator (yellow-orange areas), whereas minimum densities are achieved on the dayside close to the evening terminator (light blue areas); CO2 reaches its maximum densities in the polar region where temperatures are warmer (orange-red areas). Similarly, minimum densities of CO2 are reached at low latitudes on the nightside near morning terminator, where temperatures are colder (dark blue areas).
Fig. 4

O density maps at the (a) northern and (b) southern hemispheres; CO2 density maps at the (c) northern and (d) southern hemispheres at an altitude of 190 km. The north and south poles positions are indicated by their initials N and S, respectively. The yellow arrow points toward the sun (Valeille et al. 2009b)

2.4 Atmospheric Models

The most often used atmospheric model for Mars is the MTGCM (Bougher et al. 2008). This model solves a finite difference primitive equation that self-consistently calculates neutral, ion and electron densities over the globe under solar minimum, moderate and maximum conditions for different Mars seasons. The prognostic equations for neutral species CO2, N2, O2, O and CO are included in this model. The zonal, vertical velocity, temperature and geo-potential heights are also obtained on 33 pressure levels (above 1.32 μbar) corresponding to altitudes from 70 to 300 km with a 5 latitude and longitude resolution. The vertical coordinate is log-pressure with a vertical spacing of two grid points per scale height. Adjustable parameters that can be varied for individual MTGCM cases include f10.7 index (solar X-ray/EUV/UV flux variation), heliocentric distance (orbital variation), solar declination (seasonal variation) and maximum eddy coefficient (eddy diffusion and viscosity). The MTGCM is also modified to accommodate atmospheric inflation and semidiurnal/diurnal tidal mode amplitudes and phases consistent with dusty conditions present in Mars lower atmosphere during dust storms (Bougher et al. 2006).

The 1-D and 2-D models include the vertical and horizontal transport by eddy diffusion and convection, respectively. The 1-D thermosphere models include both chemistry and transport by eddy diffusion and do not assume a fixed homopause (e.g. Nagy et al. 1980; Krasnopolsky 2002; Fox and Yeager 2006; Fox 2009). In early 1-D models of Mars thermosphere, a single homopause altitude was assumed (e.g. McElroy 1967; Kumar and Hunten 1974; Chen et al. 1978). Below the homopause the atmosphere was considered to be completely mixed for chemical species. Above the homopause, neutral densities were assumed to be distributed according to their own scale heights. The species that are formed photochemically do not exhibit this behavior. The current 1-D models of Mars thermosphere include photoionization/excitation, photoelectron impact ionization/excitation, photodissociative excitation/ionization and more than 200 chemical reactions (e.g. Fox 2009).

3 Upper Ionosphere of Mars

Solar extreme ultraviolet radiation and precipitating particles ionize the neutral species of the Martian thermosphere producing \(\mathrm{CO}_{2}^{+}\) and O+ ions which go through the following simplified chemical scheme (Nagy et al. 2004):
$$\begin{aligned} \mathrm{CO}_2 + \mathrm{h}\nu \rightarrow& \mathrm{CO}_{2}^{+} + \mathrm{e} \\ \mathrm{CO}_{2}^{+} + \mathrm{O} \rightarrow& \mathrm{O}_{2}^{+} + \mathrm{CO} \\ \mathrm{CO} + \mathrm{O}_{2}^{+} \rightarrow& \mathrm{O}^{+} + \mathrm{CO}_{2} \\ \mathrm{O}^{+} +\mathrm{CO}_{2} \rightarrow& \mathrm{O}_{2}^{+} + \mathrm{CO} \\ \mathrm{O}_{2}^{+} + \mathrm{e} \rightarrow& \mathrm{O} + \mathrm{O} \end{aligned}$$

Initially \(\mathrm{CO}_{2}^{+}\) is produced, which is quickly removed by O leading to \(\mathrm{O}_{2}^{+}\). This ion is entirely destroyed by dissociative recombination process. The production of O+ is mostly balanced by charge exchange with CO2. This is a very fast reaction which results in the development of a major peak of \(\mathrm{O}_{2}^{+}\) at altitude ∼ 130 km during the dayside ionosphere (Nagy et al. 2004). On the nightside this peak is not well defined (Haider et al. 2009b; Lillis et al. 2011; Withers et al. 2012). The value of the peak density is determined by a balance between the production rate (qm) and the loss rate (\(\alpha N_{m}^{2}\)) at the peak altitude according to the simplified Chapman relation, \(\alpha N_{m}^{2} = q_{m} =(\eta S/eH_{n}) \operatorname{cos}(\mathrm{SZA})\), where α is the recombination coefficient (cm3 s−1), S is the ionizing flux (cm−2 s−1) at the top of the Martian atmosphere, η is the ionization efficiency (s−1), Hn is the atmospheric scale height (cm), SZA is the solar zenith angle (degree), and e=2.718. If there are no variations in Hnη, α and SZA, then Nm as well as its altitude would remain constant.

3.1 Upper Ionospheric Measurements

The ionosphere of Mars has mainly been observed by radio occultation experiments onboard ‘Mariners’, ‘Mars’, ‘Vikings’, ‘MGS’ and ‘MEX’. The radio occultation experiments onboard a spacecraft uses radio signals to probe a planetary atmosphere. The spacecraft transmits a highly stable signal which is received by the ground station. Changes in transmitted signals are attributable to change due to planetary atmospheres. As the spacecraft moves behind the planet, its radio signals cut through successively deeper layers of planetary atmospheres. Measurements of signal strength and polarization versus time yield data on the density of the atmosphere at different altitudes. It is also common to use multiple radio frequencies coherently derived from a common source to measure the dispersion of the propagation medium. This is especially useful in determining the electron content of a planetary ionosphere. In this experiment S and X band radio waves are used simultaneously to gauge the frequency dependence of the phase and amplitude variation. The S band at 2.3 GHz is sensitive to plasma density and the X band at 8.4 GHz is sensitive to the neutral density of a planetary atmospheres (Hinson et al. 1999).

3.1.1 The Early “Space Age” Measurements

The first quantitative knowledge of the upper ionosphere of Mars began with measurements made by a radio occultation experiment onboard Mariner 4 (Kliore et al. 1965) during the earliest part of the space age. Mariner 4 was followed by Mariner 6, 7 and 9 as well as by Mars 2 and 3 missions (Kolosov et al. 1972, 1973, 1975; Vasiliev et al. 1975). All these missions observed a dayside ionosphere, but there was no evidence of the nightside ionosphere (Kliore et al. 1965, 1972; Fjeldbo et al. 1966; 1970). Figure 5 shows the dayside electron density profiles observed by Mariner 4, 6 and 7 (Rasool and Stewart 1971). It should be noted that Mariner 4 measurements were made in 1965 at solar minimum, while Mariner 6 and 7 measurements were carried out in 1969 during solar maximum condition. The effect of solar activity as an increase in electron density can be noted from this figure.
Fig. 5

Electron density profiles observed by Mariner 4 (M4), Mariner 6 (M6), and Mariner 7 (M7) in the daytime ionosphere of Mars (from Rasool and Stewart 1971)

The exploration of Mars ionosphere with an orbiter started with Mariner 9, which was injected into the orbit around Mars on 14 November 1971. During the first 40 days, a total of 160 occultation measurements were performed. These data corresponded to latitudes between 34N and 65N and solar zenith angle varied from 105 to 57. Another set of occultation measurements was obtained during May-June 1972. The latitude coverage ranged from 86N to 80S with solar zenith angle varying between 70 and 100. Figure 6 shows plasma scale heights derived from all the Mariner 9 occultation measurements (Kliore et al. 1973). The average plasma scale height for the entire mission in the upper ionosphere was found to be 38.4 km.
Fig. 6

Topside plasma scale height from all data of Mariner 9 occultation measurements (from Kliore et al. 1973)

The Viking 1 and 2 spacecraft started exploring Mars ionosphere from 20 August and 9 September 1975 respectively. Most intensive investigations of the Martian ionosphere were carried out by these missions during 1976–1978 (Fjeldbo et al. 1977; Lindal et al. 1979). Radio occultation measurements were made by Viking 1 and 2 at solar zenith angles ranging from 45–127. The sub-solar and midnight ionosphere regions at Mars, which are very critical for solar wind interaction with the planet, have never been probed. Figure 7 shows the variation of peak altitudes of electron densities with solar zenith angle as observed by Mariner 9 and Viking 1/2 in the dayside ionosphere. The magnitude of the electron density maximum on the dayside depends on solar zenith angles in accordance with Chapman theory (Zhang et al. 1990).
Fig. 7

Altitude of the main ionization peak in the dayside ionosphere of Mars as a function of solar zenith angle from radio occultation data of Mariner 9, Viking 1 and Viking 2. Solid line is fitted by Chapman formula (from Zhang et al. 1990)

Like the plasmapause in the Earth’s ionosphere, a similar feature is observed in the planetary ionospheres. This feature has been given the name ionopause. Brace et al. (1983) defined it as the altitude where electron density abruptly drops to 102 cm−3. In most of the radio occultation measurements on Mariners and Vikings, there is no abrupt drop in electron density and it gradually approaches the measurement noise level. In these cases, the highest point is chosen as ionopause height. Figure 8 displays the radio occultation ionopause heights from all the Mariner 9, Viking 1 and 2 measurements (Kliore and Luhmann 1991). This figure suggests that during the period of these observations, there was no high ionopause. It should be noted that high ionopause heights often occur in the data during solar maximum and intermediate conditions for solar zenith angles greater than 55 (Ness et al. 2000).
Fig. 8

Radio occultation ionopause height plotted versus the solar zenith angle for Mars’ measurements made by Mariner 9, Viking 1 and Viking 2 (from Kliore and Luhmann 1991)

3.1.2 The Latest Measurements: MGS and MEX

During 15 years from 1965 to 1980, 443 electron density profiles were generated by radio occultation measurements on various U.S and Russian Mars’ missions (cf. Mendillo et al. 2006). Surprisingly (or unfortunately) there were no radio occultation measurements between 1980 and 1998. The radio occultation experiment on MGS observed 5600 electron density profiles during the period 24 December 1998 to 9 June 2005 (Withers et al. 2008) while a similar experiment on MEX has since reported five hundred electron density profiles (Withers et al. 2012). A number of new results have emerged from these large data sets. Each of these profiles shows a well defined peak at about 135 km for the primary layer called the F layer (Hinson et al. 1999; Pätzold et al. 2005). There is a second layer called the E layer, which often appears as a shoulder at about 112 km (Bougher et al. 2001). This is demonstrated in Fig. 9 which shows 32 electron density profiles, observed by the MGS occultation experiment during 24–31 December 1998 at different east longitudes (0–360E), small segment of latitudes (64.7–67.3N), solar zenith angles (78–81) and solar local times (0400–0300 h). It can be seen that the F peak is always present while the E peak often appears as a shoulder and sometimes as a well defined peak. Figure 10 shows a plot of F peak variations with east longitude in the northern hemisphere of Mars. Bougher et al. (2001, 2004) suggested that this variability could be created in the upper ionosphere due to the presence of nonmigrating tides/waves. Similar analyses by Seth et al. (2006a, 2006b) and Mahajan et al. (2007) confirmed the presence of wave 3 in the Mars ionosphere. Krymskii et al. (2003) have suggested that smaller and larger sized mini-magnetospheres can be formed in northern and southern hemispheres of Mars, respectively. The electron densities in the mini-magnetosphere region will be larger than those in the surrounding ionospheric regions. The higher values of peak electron densities resulted from hotter solar wind electrons that have access to the upper ionosphere along the open magnetic field lines of crustal fields (Gurnett et al. 2008).
Fig. 9

MGS radio occultation data of 32 electron density profiles for latitude 64.7 to 67.3N and SZA 78.4 to 80.8 (from Haider et al. 2006). The profiles are offset from each other by about 0.05–0.1

Fig. 10

F peak electron densities (open squares) are measured in northern hemisphere of Mars by radio occultation experiment onboard MGS. These electron densities are plotted with longitude. Solid curve indicates the best fit and dotted curves indicate 0.95 confidence limits (from Krymskii et al. 2003)

Several scientists have used radio occultation data of MGS to study the relationship of peak electron densities of E and F layers with solar zenith angles (Mendillo et al. 2003; Breus et al. 2004; Zou et al. 2006; Fox and Yeager 2006, 2009; Kumar et al. 2007; Haider et al. 2011). They have observed that the peak density and peak heights are varying significantly on the same day. Mahajan et al. (2007) referred this behavior as an anomaly since E and F layers in the Martian ionosphere are basically photochemical layers and therefore heights and densities are not expected to vary on any one day if the solar zenith angle, local time and latitude remain the same. Pätzold et al. (2005) from the radio occultation experiment onboard MEX discovered a meteoric layer between altitudes 80 km and 100 km. Withers et al. (2008) confirmed this layer from MGS radio science observations. Recently Withers et al. (2012) have reported that the electron density profiles obtained from MEX extended up to altitudes as high as 650 km in southern hemisphere in 1 % of cases. These exhibit single, double and triple scale heights in 10 %, 25 % and 10 % cases, respectively. These observations indicate that the F layer of the ionosphere can be different from its usual Chapman-like shape. These features have challenged numerical models for reproducing them and aided in understanding current ionospheric chemical, dynamical and thermal properties of Mars. Figures 11A and 11B taken from Withers et al. (2012) represent two electron density profiles illustrating variations in vertical extent of the topside ionosphere as observed by radio occultation experiment onboard MEX in northern and southern hemispheres, respectively.
Fig. 11

Two electron density profiles illustrating variations in the vertical extent of the Martian ionosphere. These profiles are measured by radio occultation experiment onboard MEX. The local time is (A) 11 h and (B) 21 h (from Withers et al. 2012)

In addition to the radio occultation experiment, MEX also carried a low-frequency radar called Mars Advanced Radar for Subsurface and Ionosphere Sounding (MARSIS) which transmits radio waves from 0.1 to 5.4 MHZ and provided topside ionograms (Gurnett et al. 2005). These ionograms provided electron density profiles in the topside ionosphere for a large range of SZA including subsolar conditions, thus providing more information on Mars morphology (Nielsen et al. 2006; Safaeinili et al. 2007; Morgan et al. 2008). The electron density profiles obtained from the MARSIS experiment have shown additional layers above the F peak altitude. Kopf et al. (2008) for example observed these layers at 280 km and 320 km on 4 September 2005 from radar sounding, as demonstrated in Fig. 12. The second and third layers have been observed in 60 % and 1 % profiles of MARSIS measurements, respectively. Another capability of MARSIS is the direct measurement of local plasma frequency (and thus the electron density) in the MEX orbit from the local plasma oscillation experiment (Picardi et al. 2004). Electron density values as low as 20 cm−3 have been measured by this technique. An important result from this experiment has been the detection of steep transient density gradients in the Mars ionosphere, similar to the Venus ionopause (Duru et al. 2009). Such sharp gradients had not been detected in any of the earlier experiments because radio occultation measurements were unable to resolve electron densities less than 103 cm−3. Figure 13 shows one such orbit showing very clearly the detection of the Mars ionopause. These sharp gradients occur somewhere between the region where ionospheric photoelectrons dominate, and the magnetosheath. The average altitude of the Mars ionopause from these measurements was found to be 500 km for SZA ≤60, after which it slightly decreases. It should be noted that sharp gradients of the type demonstrated in Fig. 13 are not observed very frequently. In 90 % of the cases, large fluctuations in electron density are seen resulting in noisy data at higher altitudes, and thus the absence of sharp gradients.
Fig. 12

Electron density profiles (circles joined by solid curve) obtained from MARSIS experiment. The red curve is a Chapman fit the maximum electron density, Ne (max). The blue and green curves are Gaussian fits to the arrowed second and third topside layers, respectively (from Kopf et al. 2008)

Fig. 13

The electron density versus UT profile for 36 min. passes on 29 September 2005. A very sharp decrease in the electron density is observed at 1523 UT at an altitude of 419 km and SZA of 28. Plasma oscillations are not observed from the beginning of the pass at 1457 UT until 1504 UT, where they suddenly appear. Similarly electron plasma oscillations suddenly disappear at about 1530 UT until the end of the pass. The sudden appearance and disappearance of the electron plasma oscillations are attributed to the high-low velocity of the shocked solar wind (from Duru et al. 2009)

3.1.3 Viking Measurements of Ion Composition and Ion Temperature

The Retarding Potential Analyser (RPA) onboard Viking 1 and 2 Landers have provided the first in situ measurements of Martian ionosphere (Hanson et al. 1977). Both Landers approached Mars from the south, crossed the equator and entered the ionosphere at solar zenith angle 45 in northern hemisphere. As stated earlier, Viking 2 and Viking 1 entered into the Martian ionosphere at 9:49 LST and 16:30 LST, respectively. The RPA experiment on Vikings 1 and 2, determined the abundance of dominant ion species and ion temperatures in the dayside ionosphere of Mars. The ion concentration profiles are shown in Figs. 14a and 14b for Viking 1 and 2, respectively (Hanson et al. 1977). It was found that \(\mathrm{O}_{2}^{+}\) is the dominant ion up to a height of about 300 km while ion O+ is likely to be dominant above this altitude. The density of \(\mathrm{CO}_{2}^{+}\) is about 10 times lower than that of \(\mathrm{O}_{2}^{+}\). The daytime ion temperature measured by RPAs (Hanson et al. 1977), departs from neutral gas temperature (TN<200 K) at altitudes well above the ionospheric peak, reaching a value of ∼3000 K near 300 km as shown in Fig. 15. Chen et al. (1978) found that EUV heating alone would result ion temperatures considerably smaller than the values measured with the RPAs. They had to assume a topside heat source flowing directly into the ions, possibly associated with some ionosphere-solar wind interaction process, to match the observed ion temperature values. Rohrbaugh et al. (1979) also carried out detailed energy calculations including contributions from exothermic chemical reactions and found good agreement between calculated and observed ion temperatures. A careful analysis of the RPA electron data allowed Hanson and Mantas (1988) to deduce the altitude variation of thermal electron temperature plus two higher energy electron populations (see Fig. 15). The first of these higher energy components is nearly Maxwellian with a temperature of Te2∼2×104 K attributable to fresh photoelectron while other higher energy population of Te3∼105 K was found to be present above 200 km possibly due to solar wind origin.
Fig. 14

(a) Concentrations of \(\mathrm{O}_{2}^{+}\), \(\mathrm{CO}_{2}^{+}\) and O+ measured by RPA on the Viking 1 Lander. (b) Same as (a) but for Viking 2 Lander observations. SZA (degree), latitude (N) and longitude (E) are given at the right as a function of altitude (from Hanson et al. 1977)

Fig. 15

Plots of temperature versus altitude. The measured electron temperature value of Te1, Te2 and Te3 were derived from least square fits, and are connected by straight lines. The measured ion temperature TI values are similarly plotted and Tn is a neutral temperature (from Nier and McElroy 1977). The short dashed curve labeled Tr is the electron temperature profile calculated by Rohrbaugh et al. (1979) and long dashed curve labeled Tc is the electron temperature profile estimated by Chen et al. (1978) (from Hanson and Mantas 1988)

3.2 Ionization Sources of Upper Ionosphere

UV and X-ray radiations are the major ionizing sources in the upper atmosphere of Mars. The photons of these frequencies contain sufficient energy to dislodge an electron from a neutral gas upon absorption. The reverse process to ionization is recombination in which a free electron is captured by a positive ion. At Mars the main reaction in the ionosphere is the dissociative recombination of \(\mathrm{O}_{2}^{+}\) and the energy is carried away in the form of kinetic energy by two resulting O atoms (see chemical reactions in Sect. 3). The ionization depends primarily on the sun and its activity. The amount of ionization in the Martian ionosphere varies greatly with the amount of radiation received from the sun. Thus there is a diurnal and seasonal effect on Mars. During the northern winter/northern summer Mars is away/close to the sun. Thus it will receive less solar UV radiation during winter than summer. The activity of the sun is also associated with the sunspot cycle, with more radiation occurring with more sunspots. Although Mars has a strong magnetic field originating in the crust, these fields are not global and therefore do not deflect solar wind particles which can ionize the upper atmosphere of Mars. However, the interaction of solar wind with inhomogeneous crustal fields are a major cause of spatial and temporal variations and thus in the ionospheric chemistry, dynamics and energetic (Nagy et al. 2004; Brain 2006; Withers 2009). Understanding the vertical structure of the Martian ionosphere has been a key tool for discovering how the ionosphere behaves and what processes shape it (cf. Mendillo et al. 2011). The upper ionosphere of Mars is divided into E and F region. These two regions are described in the following sections.

3.2.1 Ionization by X-Rays: E Region Ionosphere

Mars ionospheric E layer, which has its peak density at about 115 km, is produced by X-ray radiation with wavelengths of 10–90 Å. The E regions of Mars and Earth’s ionosphere have been compared by several investigators (Haider et al. 2002; Mendillo et al. 2003; Rishbeth and Mendillo 2004; Haider et al. 2009a, 2009b). They have found that the E layer is created at nearly the same height in the ionospheres of both planets, but the layer thickness is considerably less on Mars than that on Earth. At night the E layer disappears because the primary source of ionization is no longer present. The vertical structure of the E layer is primarily determined by the competing effects of ionization and recombination. The ion composition calculation shows that \(\mathrm{O}_{2}^{+}\) and NO+ are the major ions in the E region with \(\mathrm{N}_{2}^{+}\) and O+ as minor ions (Haider et al. 2012). Within the E region of Earth’s ionosphere, local enhancements and irregularities in the electron density, known as sporadic, are observed (e.g. Rishbeth and Garriott 1969; Hargreaves 1992). Sporadic—E events may last for few minutes to several hours. These events occur mostly in the summer season when high solar insolation reaches on Earth. Figure 16 shows production rate (photoionization + photoelectron impact ionization rates) on 29 May, 2003 for major ions \(\mathrm{CO}_{2}^{+}\), \(\mathrm{N}_{2}^{+}\) and O+ in the E region of Martian ionosphere due to absorption of solar X-ray radiation during quiet condition. X-rays produce major ion \(\mathrm{CO}_{2}^{+}\) in the E region of the Martian ionosphere and the maximum ion production occurs in the altitude range ∼ 90–110 km (Haider et al. 2012).
Fig. 16

Altitude profile of predicted ion production rates of \(\mathrm{CO}_{2}^{+}\), \(\mathrm{N}_{2}^{+}\) and O+ in the E region of Martian ionosphere due to absorption of X-ray radiation (1–8 Å) during quiet conditions (from Haider et al. 2012)

3.2.2 Ionization by Solar EUV: F Region Ionosphere

The F region is formed mainly owing to photoionization of neutral atoms/molecules by EUV radiation from 90–1026 Å. Precipitating energetic charged particles also contribute to ionization of neutral species. The major gases CO2, N2, O2, O, Ar and CO are ionized by EUV radiations. The maximum ionization in the F region occurs at altitude ∼125–135 km in the dayside ionosphere of Mars for sub-solar conditions (cf. Hanson et al. 1977; Fox and Dalgarno 1979; Bougher et al. 1990; Fox 2009). At higher solar zenith angles of the dayside, the peak ionization occurs at higher altitudes. Recently Haider et al. (2009b, 2010a) have compared the ionospheric F regions of Earth and Mars. They have found that the thickness of the F layer and location of its peak density decreases by a factor of 1.6 to 1.8 in the Martian ionosphere as compared to that observed in the Earth’s ionosphere because of smaller neutral scale height, the solar EUV energy is deposited within smaller altitude range in the upper ionosphere of Mars as compared to the corresponding altitude range in the upper ionosphere of Earth. Figures 17 and 18 show sample profiles of electron density observed at nearly the same time, location and solar conditions in the ionospheres of Earth and Mars, respectively (Haider et al. 2009b). Six profiles have been selected in the Earth’s ionosphere from COSMIC measurements carried out on 20 March 2008 at coordinates (72.7N, 153.4W), (51.8N, 49.5W), (49.8N, 144.5W), (67.3N, 146W), (54N, 126.3W) and on 9 April 2008 at coordinates (58.4N, 93.6W). These observations were carried out in the afternoon at nearly the same local time (13.2, 13.6, 13.8, 14.4, 13.2 and 14.3) during a low solar activity period F10.7=68). Six profiles of electron density have been selected from MGS observations carried out on 20 March 2005 at coordinates (74.4N, 106.5W), (74.3N, 49.2W), (74.2N, 209W), (70N, 156W), (70N, 127W), and on 9 April 2005 at coordinates (70N, 99W) (The data are not available in 2008 because MGS failed to work after 9 June 2006). These observations were again performed in the afternoon at nearly the same local times (12.95, 12.96, 12.97, 13.75, 13.76 and 13.74) with low solar activity (equivalent F10.7=39 at Mars).
Fig. 17

Six sample profiles of electron density representing D, E and F layers as observed in Earth’s ionosphere by radio occultation experiment onboard COSMIC satellites at different locations at nearly the same local time in the afternoon (from Haider et al. 2009b)

Fig. 18

Six sample profile of electron density representing E and F layers in the Martian ionosphere as observed by radio occultation experiment onboard MGS at nearly the same local time in the afternoon. Density profiles representing D layer between altitude range ∼25–35 km is observed by radio occultation experiment onboard Mars 4 and Mars 5 at 04:30 LST (from Haider et al. 2009b)

3.2.3 Solar Wind Impact Ionization

Outside the crustal magnetic field region Mars has an induced magnetosphere. As a result solar wind dynamic pressure compresses the interplanetary magnetic field into the Martian ionosphere. Shinagawa and Cravens (1989) developed a one-dimensional magnetohydrodynamic (MHD) model to study the role of electromagnetic forces in the Martian ionosphere. They found a good agreement between model and Viking observations by adding an extra heat source caused by the solar wind in the upper ionosphere. Later Shinagawa and Bougher (1999) developed a two-dimensional MHD model and studied two cases of solar wind dynamic pressure on Mars. In the first case solar wind velocity was high ∼450 km/s, which exceeded the maximum ionospheric pressure of Mars. In the second case, solar wind velocity was below ∼300 km/s. This study showed that the upper ionosphere of Mars is significantly influenced by solar wind. Using a three-dimensional model, Ma et al. (2004) have reported that solar wind plays an important role above 250 km. Similar conclusions were also drawn by Haider et al. (2010a) using a one-dimensional model with non-zero upward flux boundary condition. Below this height, these models are in agreement with the earlier modeling results.

The nightside ionosphere of Mars was first measured by the radio occultation experiment onboard Mars 4 and Mars 5 at solar zenith angles of 127 and 106 in February 1974 under solar minimum condition (Savich and Samovol 1976). Later Zhang et al. (1990) reported that about 60 % of the radio occultation profiles obtained in the nighttime from Viking 1/2 do not show a well defined peak during low solar activity. For the remaining 40 % profiles the average nightside peak value was about 5×103 cm−3, at an altitude of about 150 km. This peak is produced by solar wind electron transportation from dayside to nightside atmosphere across the terminator (cf. Verigin et al. 1991; Haider et al. 1992; Fox et al. 1993). Kallio and Janhunen (2001) have calculated ion production rates in the nightside Martian ionosphere due to H+–H impact process. This source was found to be an important ionization process for the nightside ionosphere of Mars. Using this source of ionization, Haider et al. (2002) predicted peak electron densities of 3.5×103 cm−3 and 2.0×103 cm−3 at solar zenith angles of 105 and 127 respectively. They found that fast hydrogen atoms penetrate deeper into Martian atmosphere and lose their energy at lower altitudes as compared to solar wind electron impact ionizations. Above 200 km, the photoelectrons produced during the dayside that travel to the nightside are found to be an important process that contributes about 30–40 % photoelectron flux in the nighttime ionosphere for SZA ∼127 (Fox et al. 1993).

3.3 Chemistry of the Upper Ionosphere

The chemistry of ions \(\mathrm{O}_{2}^{+}\), NO+, \(\mathrm{CO}_{2}^{+}\), O+, \(\mathrm{N}_{2}^{+}\), and CO+ in the upper ionosphere of Mars has been studied by several investigators (cf. Chen et al. 1978; Hanson et al. 1977; Fox 1993; Haider 1997; Ma et al. 2004; Duru et al. 2008; Haider et al. 2010a; Haider 2012). Solar EUV is the major ionization source for the production of these ions in the dayside while it is electron impact ionization for the nightside ionosphere of Mars. Major ions are \(\mathrm{O}_{2}^{+}\), NO+ and \(\mathrm{CO}_{2}^{+}\) below about 200 km above which O+ dominates. The ion NO+ is mainly produced due to reaction of \(\mathrm{O}_{2}^{+}\) with N and NO, and it is entirely destroyed by dissociative recombination. The density of NO+ is directly proportional to the densities of N and NO (Haider et al. 2009b). The other major source of this ion is the reaction of CO+ and \(\mathrm{N}_{2}^{+}\) with CO2. The ion \(\mathrm{CO}_{2}^{+}\) is lost by collision with atomic oxygen and is the one of the main sources of the dominant ion \(\mathrm{O}_{2}^{+}\). Dissociative recombination is an important loss process for \(\mathrm{O}_{2}^{+}\) at all altitudes. The dominant ion \(\mathrm{O}_{2}^{+}\) is produced in the dayside and nightside ionosphere mainly due to the reaction of \(\mathrm{CO}_{2}^{+}\) with O (Barth 1985). Among the minor ions, CO+ is lost in the charge exchange reaction with CO2. This process destroys almost all CO+ ions in the Martian ionosphere. Solar wind electron impact ionization is the major source of CO+ in the nightside ionosphere (Haider et al. 2007). Fox (2009) found that the charge exchange reaction between \(\mathrm{CO}_{2}^{+}\) and O is the dominant source of O+ at heights below 190 km. Above this altitude solar radiation is an important production mechanism for O+. The ion \(\mathrm{N}_{2}^{+}\) is produced by photoionization and is lost by reaction with CO2. The chemical reaction scheme for the upper ionosphere of Mars is shown in Fig. 19.
Fig. 19

Ionospheric Chemical reaction scheme of Mars upper ionosphere (from Chen et al. 1978). This scheme has been used and updated by various workers (references are given in this review paper)

3.4 Ionospheric Models

For studies of ionospheric physics, Boltzmann equation is represented by continuity, momentum and energy equations. The continuity equation is written as:
$$ \frac{\partial n_{s}}{\partial t} + \nabla \cdot (n_{s}\mathbf{u}_{s}) = P_{s} - L_{s} $$
where Ps is the production rate of species s including primary production by photoionization and photoelectron impact ionization (or by collisional ionization by energetic electrons or ions) as well as the production due to chemical reactions and Ls is the loss rate of species s due to chemistry of various reactions. This equation has been used by many investigators to study the ionosphere of Mars (e.g. Fox 1993, 2009, Fox and Yeager 2006, 2009; Haider et al. 2010a and references therein). The momentum equation for species s can be written as:
$$\begin{aligned} n_{s}m_{s} \biggl[ \frac{\partial \mathbf{u}_{s}}{\partial t} + \mathbf{u}_{s} \cdot \nabla \mathbf{u}_{s} \biggr] =& - \nabla p_{s} + n_{s}e_{s} ( \mathbf{E} + \mathbf{u}_{s} \times \mathbf{B} ) + n_{s}m_{s} \mathbf{g} \\ &{} - n_{s}m_{s}\sum_{j} v_{sj} ( \mathbf{u}_{s} - \mathbf{u}_{j} ) + p_{s}m_{s} (\mathbf{u}_{s} - \mathbf{u}_{n} ) \end{aligned}$$
where g is the acceleration due to gravity and νsj is the momentum transfer collision frequency between species s and j. The charge on species s is equal to ±e depending on whether s denotes electrons or ions, ms is the mass of charged particles, E and B are the electric and magnetic fields, respectively, ps is the pressure (=nskBTs where kB and Ts are Boltzmann constant and temperature, respectively) and un is the speed of the neutral particle. The temperature for species s can be calculated by the energy equation
$$ \frac{3}{2}k_{B}n_{s}\frac{\partial T_{s}}{\partial t} - \frac{\partial}{ \partial z} \biggl( K_{s} \frac{\partial T_{s}}{\partial z} \biggr) = Q_{s} - L_{s} $$
where Qs and Ls are heating and cooling rates, respectively and Ks is coefficient of thermal conductivity. This energy equation leaves out dynamical terms such as heat advection, which is less important for electrons than heat conduction and local heating and cooling. Chen et al. (1978) and Rohrbaugh et al. (1979) have solved energy equations to study ion and electron temperatures in the dayside ionosphere of Mars. Their results were compared with the temperature measurements made by Viking 1 in Fig. 15. In the magnetized plasma, the momentum equation includes the magnetic field (cf., Eq. (2)), which can be derived from the so called magnetic induction equation. The induction equation can be written as:
$$ \frac{\partial \mathbf{B}}{\partial t} = \nabla \times ( \mathbf{u}_{s} \times \mathbf{B} ) - \nabla \times \biggl( \frac{\eta}{ \mu_{o}} \nabla \times \mathbf{B} \biggr) $$
This equation is obtained by combining Faraday’s law, the generalized Ohm’s law and Ampere’s law and is known as the magnetic convection-diffusion equation. The first term on the right hand side of Eq. (4) is magnetic convection term, and the second term is the magnetic diffusion term.

Apart from continuity, energy and momentum equations, there are MHD and hybrid models, which are self-consistent Mars-solar wind plasma approaches (cf. Nagy et al. 2004; Ma et al. 2008; Kallio et al. 2010). MHD model provides a high-resolution, three-dimensional simulation of the Martian ionosphere, which contains both a solar wind and a self consistent ionosphere. In this model all ions are assumed to have the same bulk velocity. Furthermore, it assumes a Maxwellian velocity distribution function, while pickup ions O+ are highly non-Maxwellian on Mars. The hybrid model is also a three- dimensional model and has been used to study the global modeling of the Mars-solar wind interaction (Modolo et al. 2006; Brecht and Ledvina 2006; Ledvina et al. 2008; Brain et al. 2010; Kallio et al. 2010). This model represents electrons as a massless fluid. The advantage of this model is that it includes kinetic effects such as finite gyro radius of ions, wave particle interaction and instabilities associated with the non-Maxwellian velocity distribution function. Hybrid model has a low resolution. It does not include Martian crustal magnetic anomalies because the limited resolution makes it impossible to include a realistic crustal magnetic field model.

4 Mars’ Upper Ionospheric Disturbances

No significant planetary magnetic field of global scale has been detected at Mars. The magnetometer experiment onboard MGS has reported the upper limit of Mars’ dipole moment ∼2×1022 Gauss-cm3, which is no more than 10−4 times as strong as that of the Earth (Acuña et al. 1998). Analysis of meteorites show that Mars used to have a strong planetary magnetic field but lost it ∼4 billion years ago (Acuña et al. 1998; Lillis et al. 2008; Weiss et al. 2008). Thus for most of Mars’ history, it has had induced magnetosphere and solar wind interaction. Figure 20 shows a global map of the crustal magnetic field amplitude over Mars at an altitude ∼170 km (Mitchell et al. 2007). The crustal fields are stronger in the southern hemisphere than in the northern hemisphere.
Fig. 20

Global map of the crustal magnetic field amplitude (B) of Mars at altitude 170 km as inferred from electron reflectometry represented by colors superposed on shaded relief map of MGS Mars Orbiter Laser Altimeter topography. Labels indicate the centers of the Utopia (U), Hellas (H), Argyre (A), and Isidis (I) impact basins, as well as the volcanoes Elysium (e), Syrtis Major (s), Applinaris (o), Tyrrhena (t), Hadriaca (h), Peneus (p), and Amphitrites (a), Olympus Mons is represented by the thick black region, where magnetic field lines are shown from closed loops (Mitchell et al. 2007; Haider et al. 2010a)

The variations in the characteristics of solar wind can induce a series of disturbances in the upper ionosphere of Mars. Further, when a large flare occurs on the solar surface, a vast amount of X-rays, high energy particles and plasma masses are emitted from the active region of sun into the interplanetary space. An encounter of Mars with these energetic streams results in violent disturbances of the upper atmosphere. In the following section we will review aurora, magnetic storm and ionospheric perturbations due to impact of solar flares, CMEs and SEPs on the upper atmosphere of Mars.

4.1 Solar Flares and Upper Ionospheric Disturbances

Several solar flares of classes M and X were registered on the Sun during the maximum phase of solar cycle 23. Responses to these flares in the Martian ionosphere have been reported by several investigators. Using MARSIS data, Nielsen et al. (2006) reported that the maximum electron density in the Martian ionosphere suddenly increased from 1.8×105 to 2.4×105 cm−3 at 08:39 UT on 15 September 2005. This coincided closely in time with the increase in solar X-ray fluxes measured from GOES 12 spacecraft at the Earth, as shown in Fig. 21. Mendillo et al. (2006) examined ionospheric data of MGS and reported two elevated electron density profiles during the solar X-ray flares which occurred on 15 April and 26 April 2001 at 13:50 UT and 13:10 UT respectively. They found 200 % enhancement in the electron density profiles during these flares, as shown in Fig. 22. Haider et al. (2009a) studied solar flare of 13 May 2005 and the following CMEs and reported their effects in the E region ionosphere. Mahajan et al. (2009) surveyed all the electron density profiles measured by MGS and reported effects of seven X-ray flares in the Martian ionosphere. They found elevated electron densities in the E region during all the flares but in the F region during some flares only.
Fig. 21

Sudden increase of electron density maximum (top panel) simultaneous to an increase in solar X-ray flux (bottom panel) of wavelength 0.5–4 Å (lower curve) and 1–8 Å (top curve). The X-ray data were observed on September 15, 2005 by GOES spacecraft when MARSIS observed the sudden increase in electron density (from Nielsen et al. 2006)

Fig. 22

(A) Electron density profiles on Mars obtained for 15 April and 26 April, 2001. Two profiles in red at 14:15 and 13:16 UT show significant enhancement at low altitude because of solar flares which peaked in X-ray fluxes at Earth at 13:50 and 13:10 UT respectively. On 15 April there were five MGS profiles before the flare at 2:28, 6:23, 8:21, 10:19 and 12:17 UT and none after the flare. On 26 April pre flare profiles were available at 9:20 and 11:18 UT and post flare at 17:11 and 19:09 UT. (B) % differences between the flare affected profiles and the averages of the other profiles on each day. The shading gives 1−σ standard error in relative change in electron density (Ne) (from Mendillo et al. 2006)

Recently Haider et al. (2012) have modelled solar X-ray fluxes measured by GOES 12 during the periods 29 May to 3 June 2003, 15–20 January 2005 and 12–18 May 2005 and investigated effects on electron densities produced by individual X-ray flares that occurred within these intervals. They have reproduced the major characteristics and magnitudes of the measured solar X-ray spectra by model calculation. Figures 23a, 23c and 23e present the calculated and measured solar X-ray flux distribution from 29 May to 3 June 2003, 15 to 20 January 2005 and 12 to 18 May 2005 respectively in the wavelength range 1–8 Å. Figures 23b, 23d and 23f show the distribution of proton fluxes with time as observed by GOES 11 at three energies ≥10 MeV, ≥ 50 MeV and ≥ 100 MeV thus indicating that CMEs erupted from the active regions on the Sun for several hours after each flare. The flare of January 17 was very strong, and protons of all three energies were accelerated from the heliosphere of the Sun (Fig. 23d). The X-ray and proton fluxes for the flares of 29 May and 31 May 2003 were not as large as those for the flares of 17 January and 13 May 2005. Figures 24a, 24b and 24c represent time series of the measured and calculated Total Electron Content (TEC) of E-region ionosphere for each respective day between 29 May and 3 June 2003, 15 and 20 January 2005 and 12 and 18 May 2005 respectively. MGS observed the electron density profiles a few hours before and immediately after these solar flares. Before flaring, the ionosphere of Mars was calm. Soon after the solar flares, an increase by a factor of ∼ 4–5 in the TEC was estimated. The effect of these flares endured each day for about 1–2 h in the E region ionosphere of Mars. During the quiet period the value of predicted TEC is higher by a factor of 1.5 to 2 than the observed value. This is due to the fact that electron-ambient-electron collisions were neglected in this model. This process will reduce the modelled TEC.
Fig. 23

(a) Solar X-ray flux measured by GOES 12 (black line) and calculated by model (red line) between 29 May and 3 June 2003. (b) Proton flux distributions between 29 May and 3 June 2003 measured by GOES 11 at three energies: ≥10 MeV (red color), ≥50 MeV (black color) and ≥100 MeV (blue color). (c) Solar X-ray flux measured by GOES 12 (black line) and calculated by model (red line) between 15 and 20 January 2005. (d) Proton flux distribution between 15 and 20 January 2005 measured by GOES 11 at three energies: ≥10 MeV (red color), ≥50 MeV (black color), and ≥100 MeV (blue color). (e) Solar X-ray flux measured by GOES 12 (black line) and calculated by model (red line) between 12 and 18 May 2005. (f) Proton flux distribution between 12 and 18 May 2005 measured by GOES 11 at three energies: ≥10 MeV (red color), ≥50 MeV (black color), and ≥100 MeV (blue color) (from Haider et al. 2012)

Fig. 24

(a) Measured TEC (blue color) and predicted TEC (red color) in the E region ionosphere of Mars for the period 29 May to 3 June 2003. (b) Measured TEC (blue color) and predicted TEC (red color) in the E region ionosphere of Mars for the period 15–20 January 2005. (c) Measured TEC (blue color) and predicted TEC (red color) in the E region ionosphere of Mars for the period 12–18 May 2005. CME arrival and its effect on 30–31 May 2003, 2–3 June 2003 and 16–17 May 2005 are marked by an arrow line (from Haider et al. 2012)

4.2 Effect of CMEs on the Upper Ionosphere

Studying and understanding the effect of CME is a key area of research in planetary aeronomy. Haider et al. (2012) were the first to detect the effect of CMEs in the E region ionosphere of Mars. They examined MGS data between 30 and 31 May 2003, 2 and 3 June 2003 and 16 and 17 May 2005 following the flares of 29 May, 31 May 2003 and 13 May 2005 respectively (see Figs. 24a–24c). They found that the physical processes of magnetic storms (the after effects of CMEs) are different on Earth and Mars. During a magnetic storm, shock waves driven by CME compress the Earth’s magnetosphere leading to increased energetic particle precipitation into the ionosphere. This leads to sudden increase in the electron density. The magnetic storm effects on Mars, on the other hand, have quiet different characteristics. Mars has no dipolar magnetic field. Therefore solar wind interacts directly with the Martian ionosphere, which acts as an obstacle and diverts the solar wind around it (Acuña et al. 1998). Actually it is a magnetic barrier, which diverts the bulk of solar wind shock around the planet. There exists a sharp boundary (ionopause) between the magnetised magnetosheath and the ionosphere, as shown in Fig. 25. The location of the magnetosheath and ionopause can change with the solar wind dynamic pressure. Thermal pressure balances magnetic pressure at the ionopause boundary. Mass loading of the magnetosheath does effect the flow behind the shock such that the pickup of neutrals ionized outside the ionopause (i.e. pick up ions) contribute to further stagnation of this flow and to the growth of the magnetic barrier. Consequently this contributes to the development of the magnetotail.
Fig. 25

Schematic diagram of the ionopause between magnetized magnetosheath and the ionosphere of Mars. Usw and Bsw represent the velocity and magnetic field directions of solar wind on Mars respectively. Magnetic barrier, bowshock and ionopause locations are also shown

MGS has observed a magnetosheath at about 435 km on the sunlit hemisphere of Mars during quiet conditions (Mitchell et al. 2000). In the magnetosheath the planetary neutrals are mainly H atoms of the hydrogen corona. Fast hydrogen atoms are produced by charge exchange between solar wind protons and hydrogen corona in this region. These energetic proton-hydrogen atoms have the same energies as the solar wind protons and move in the same direction as that of the fast protons just before the collisions (Haider et al. 2002). In this way the magnetosheath of Mars can be compressed similar to that observed in the Earth’s magnetosphere (Dandouras et al. 2007) and the accelerated solar wind protons get turned into fast hydrogen atoms at lower altitudes. To verify that the flare of 13 May 2005 has an effect on the magnetosheath of Mars, Haider (2012) analyzed the magnetic field data obtained from MGS at altitudes ∼ 420 km and ∼ 430 km from 12 to 18 May 2005. The variation of magnetic field measured in the magnetosheath region of Mars is shown in Fig. 26. There are two broad peaks in the magnetic field at an altitude of ∼420 km on 15 and 17 May with the values ∼50 nT and 40 nT at 21:50 UT and 02:52 UT respectively. These values are larger than that the magnetic field normally observed at altitude ∼430 km by a factor of ∼2.5. Before and after these times, the magnetic field does not change significantly between these two altitudes. This suggests that the CME arrived at Mars on 15 May at about 21:50 UT and compressed its magnetosheath by about 10 to 15 km.
Fig. 26

Magnetic field in the magnetosheath region of Mars at altitude 420±2 km (red line) and 430±5 km (blue line) passing through MGS at latitude 65–66N between 12 and 18 May 2005. The major peaks at 21:50 UT and 02:52 confirms the arrival of CME at Mars on 15 and 16 May, respectively. Dashed blue line shows the magnetic field magnetosheath of Mars under quiet condition (from Haider 2012)

Haider et al. (2012) also ran a three-dimensional kinetic solar wind model (Hakamada-Akasofu-Fry Version 2/HAFv.2) to confirm the arrival of CME at Mars following the solar flare of 13 May 2005. This model does not provide any way to distinguish between the effects on the ionosphere of magnetic storm from the CME shocks and the energetic particles from that shock. Figures 27a–27h shows simulated ecliptic plane profile of IMF (about to 2 AU) from 15 to 18 May 2005. The simulation confirmed that the CME reached Mars on 15 May 2005 after its time of arrival at Earth (Figs. 27b and 27c). A second CME-related shock was predicted to reach Mars on 17 May 2005 at 16:00 UT (Fig. 27g). The arrival of the resultant shock was predicted to reach Mars between 16 and 17 May 2005 (Figs. 27d–27g). The direction of the CME was away from Mars on 18 May (Fig. 27h). The overall effect of CME arrival at Mars lasted for about 2 days. As a result TEC increased suddenly (see Fig. 24c) by factors of ∼2–3 between 16 and 17 May 2005.
Fig. 27

Ecliptic plane simulations out to 2 AU of IMF and solar wind disturbances predicted by the HAFv.2 model during 15–18 May 2005 at selected time: (a) 15 May at 0.00 UT, (b) 15 May at 8:00 UT, (c) 15 May at 16:00 UT, (d) 16 May at 0:00 UT, (e) 16 May at 12:00 UT, (f) 17 May at 0:00 UT, (g) 17 May at 16:00 UT and (h) 18 May at 0:00 UT. IMF lines are shown in red (away sectors) and in blue (towards sectors). The locations of the Earth and Mars are indicated by black dots (from Haider et al. 2012)

4.3 Effect of SEPs on the Upper Ionosphere

Solar Energetic Particles (SEPs) events are a part of major disturbances in the heliosphere (Schnjver and Siscoe 2010). These events are mostly composed of protons with about 10 % He+ and <1 % heavier elements. There are two types of SEP events: Impulsive and gradual (Cane et al. 1986). Impulsive events are relatively of short duration (<1 day) with a high proton content. Gradual events are of longer duration (days), have higher fluxes, display a wider spread in longitude and are associated with fast CMEs. McKenna-Lawlor et al. (2012) reported three major factors in connection with effect of SEP radiations at the Martian surface: (1) shadowing by the planet, which cuts off ∼50 % of SEP primary particle flux, (2) atmospheric attenuation, which shields out SEP primaries characterised by relatively low energies and (3) backscattering particles, mostly neutrons, due to the interaction of high energy particles with soil material. The Mars Energetic Radiation Environment models (MEREM) were developed by ESA and NASA to study the SEP radiation close to the Martian environment. The output of these models gives (1) particle influence, (2) effective dose and (3) ambient dose equivalent with the Martian atmosphere and in the planetary orbit about Mars (McKenna-Lawlor et al. 2012) It is known that the energetic particle densities in the solar wind are significantly enhanced during SEP events. The observations made by MARSIS have demonstrated that SEP events modify the ionosphere of Mars (Morgan et al. 2006). It needs to be mentioned that this instrument did not detect its usual reflections from the surface of Mars during an SEP event, indicating that radio waves which usually pass smoothly through the ionosphere were fully absorbed (Withers 2011). Sheel et al. (2012) investigated the effect of SEP event of 29 September 1989 on the ionosphere of Mars. This event was observed by IMP and GOES satellites during the disturbed condition of the Sun (Lovell et al. 1998). Figure 28 shows electron density profiles predicted for the 29 September 1989 SEP event, with and without photoionization process. It can be noted that the SEP event caused electron density to exceed 2×105 cm−3 at ∼120–140 km, a value much larger than typically observed by MGS in the dayside ionosphere of Mars (cf. Haider et al. 2011). Sheel et al. had carried out two model calculations—photochemical and generalized models. Solid lines represent the calculation of photochemical equilibrium model. The black dashed line shows the electron density profile obtained from a generalized model. The generalized model neglects the energy deposition at high energies. Therefore, electron density decreases exponentially with height in this model. The photochemical equilibrium model includes photoionization process, chemical reactions and their production and loss processes under steady state condition. There is a reasonable agreement between two model calculations in the absence of photoionization. Before MARSIS observations, Leblanc et al. (2002) had simulated the vertical profile of energy deposition rate in the Martian ionosphere for the SEP event of 20 October, 1995. However, they did not study the ionospheric effects of this SEP event at Mars. Their results are comparable with Sheel et al. (2012) in the vicinity of ionization peaks produced by this SEP event.
Fig. 28

Electron density profile predicted for 29 September 1989 SEP event, neglecting photoionization (grey solid line) and including photoionization (black solid line). These two profiles are identical below 100 km. The black dashed line shows the electron density profile from generalized model (from Sheel et al. 2012)

4.4 Auroral Ionosphere

Aurorae are often seen in the upper atmosphere of Earth at high latitudes following the occurrence of solar flares and CMEs. Since Mars has no strong remnant of an ancient intrinsic magnetic field in the northern hemisphere (Acuña et al. 1998), Fox (1992) argued that Mars should have Venus-like diffuse aurora in this region. However, auroral events are not observed in the northern hemisphere, but effects of magnetic storms have been detected in the ionosphere of Mars (Haider et al. 2009a; Haider 2012). In Sect. 4.2, we explained that the physical processes of Earth’s and Mars’ aurorae are different. Bertaux et al. (2005) have discovered southern aurora in the nighttime atmosphere (solar zenith angle 117.5) of Mars using the “Spectroscopy for the Investigation of the Characteristic of the Atmosphere of Mars” (SPICAM) experiment onboard MEX. Figure 29a shows the limb observations carried for 450 to 750 seconds at wavelength range of 100–350 Å in the nighttime atmosphere. H lyman α (121.6 nm) and NO bands (181–298 nm) are clearly visible in this figure. Figure 29b represents an auroral spectrum integrated over the wavelength range of the NO bands (181–298 nm) as a function of time for the five spatial bins. There is a strong peak in all the bins. These spectra were observed at a tangent altitude of 19 km. Local time was ∼21:00 h and longitude and latitude were 198.4 and −46.3 respectively. These measurements were carried out using nadir looking direction in the strong crustal magnetic field region where the field lines are nearly open (Lundin et al. 2006; Mitchell et al. 2007; Haider et al. 2010a). Leblanc et al. (2008) found a good correlation among the measured auroral emission by SPICAM, the measured downward/or upward flux of electrons by ASPERA-3 and the TEC recorded by MARSIS. They found that TEC increased when there was a precipitation of high flux auroral electrons into the atmosphere of Mars. During the auroral events ASPERA-3 measured an increase in the electron flux by an order of magnitude. Before and after this event the electron flux decreased by two orders of magnitude. Leblanc et al. too found a broader peak in TEC during the SPICAM measurements of an auroral event. This suggests that electron density increased significantly in the auroral ionosphere of Mars during this event. Figures 30a–30g displays the measurements by SPICAM (Fig. 30g), MGS/Electron Reflectometer (ER) (Fig. 30f), ASPERA-3 (Figs. 30d and 30e), MARSIS (Fig. 30c) and the MEX altitude (Fig. 30a) and MEX latitude (Fig. 30b). MEX observations were carried out on 26 January 2006 during orbit # 2621. Track of MEX during this orbit was at local time of 20:30 h at a longitude of 180 that is above the most intense crustal magnetic field recorded by MGS (Mitchell et al. 2007; Haider et al. 2010a). SPICAM observed a significant increase of auroral light between 14:04:01 UT and 14:04:15 UT.
Fig. 29

(a) Spectra recorded during the grazing limb observation between 450 to 750 seconds. Altitudes of the Mars Nearest Point (MNP) of the line of sight are indicated at the right. It contains H Lyman α emission at 121.6 nm and well structured band (190–270 nm) of NO. The intensity in ADU (Analogue Digital Units) per pixel is color-coded. (b) Auroral peak is sharp and different from NO spectrum. Signal intensity in ADU for five bins (each averaged from 181–298 nm) as a function of time between 200 and 900 seconds are plotted in Fig. 32b (from Bertaux et al. 2005)

Fig. 30

Time series of MEX measurements during orbit 2621. Shown is the MEX (a) altitude, (b) latitude, (c) MEX/MARSIS TEC, (d) MEX/ASPERA-3/ELS electron measurement, (e) MEX/ASPERA-3/ELS energy flux, (f) MGS/MAG/ER electron energy flux and (g) MEX/SPICAM measurements. Indicated on each by vertical dotted lines are the periods during which an aurora event has been identified in SPICAM/UVS observations (from Leblanc et al. 2008)

It is found that the energy distribution of the downward electron flux measured by ASPERA-3 exhibited non-Maxwellian features similar to those observed in the V-shaped potential structure of Earth’s auroral zone (cf. Fang et al. 2010; Lei and Zhang 2009; Lillis et al. 2009; Fillingim et al. 2010; Lundin et al. 2011). Ip (2012) proposed a mechanism for the acceleration of ions and electrons in the Martian auroral zone. Figure 31 represents a schematic view of inverted V shaped potential structures, where electrons are precipitating downward and conic ions are escaping upward in the presence of strong crustal magnetic fields. Ip suggested that only those ions, which are created in the crustal fields connected to the interplanetary magnetic fields, would be able to escape. The ions which are created in the magnetic flux tubes of closed field lines will be trapped in a bouncing motion. The southern aurora may not be visible to the human eyes because they were observed in the ultraviolet wavelength region.
Fig. 31

Schematic view of the downward electron beams and the formation of upward flux of oxygen ions at the foot points of the Martian crustal magnetic fields. Transverse heating of the O+ ions and other ion species will lead to the generation of ion conics, which are to be converted to ion beams at higher altitude. Parallel electric fields will be maintained by the V-shaped potential structure (from Ip 2012)

4.5 Modeling of the Upper Ionosphere During Disturbed Condition

Fox et al. (1996) carried out a model calculation of the thermosphere and ionosphere of Mars and calculated neutral densities, temperature, ion production rates and densities of ions and electrons during solar minimum and maximum conditions. They found that the electron density increased by a factor of ∼3 during solar maximum condition in comparison to that estimated for solar minimum condition. These model results were found to be in good agreement with the radio occultation measurements made by Viking 1 and Mariners 6 and 7 during solar minimum and solar maximum conditions respectively. Fox et al. (1996) used Eq. (1) to calculate the ion and electron densities in the Martian ionosphere. In this calculation the model atmospheres for low and high solar activities were used from the MTGCM model of Bougher et al. (1990). Using the photochemical model, Mendillo et al. (2004) estimated noontime TEC at sub-solar latitude of Mars for aphelion and perihelion positions during solar minimum and solar maximum periods. In this calculation neutral atmosphere and solar flux were taken from Bougher et al. (1990) and Tobiska et al. (2000) respectively for solar maximum/minimum conditions. The results of these model calculations are shown in Fig. 32. It can be seen that TEC increased in the Mars’ ionosphere by a factor of ∼2 during solar maximum conditions. It should be noted that TEC in the Earth’s ionosphere is larger during solar maximum conditions by one to two orders of magnitude (Rishbeth and Mendillo 2001) not because of Earth’s closer distance to the sun, but due to the non-photochemical layer F2, providing major contribution to the TEC integral. Thus photochemical models in Mars’ ionosphere exhibit minimum solar cycle variations.
Fig. 32

Model results for noontime ionospheric TEC in the altitude range 100–200 km at the sub-solar latitude on Mars for aphelion/perihelion positions during solar maximum/minimum conditions (from Mendillo et al. 2004)

Using MARSIS data obtained from MEX, Lillis et al. (2010) have reported that disturbed solar and space weather conditions can produce prolonged higher TEC values, while an individual SEP event causes a short-lived absolute increase in TEC. They also found a relationship between TEC and both He-II line irradiance and F10.7 solar radio flux as power laws with exponents of 0.54 and 0.44 respectively. As mentioned before, Haider et al. (2012) used radio occultation data of MGS at high latitudes (65.3–65.6N, 69.3–69.6N and 74.6–77.5N) and studied the effects of solar X-ray flares on TEC in the E region of the Martian ionosphere. Modeling of flare-induced solar X-ray fluxes, ion production rates, electron densities and TEC were carried out for solar flare events that occurred on 29 and 31 May 2003 and 17 January and 13 May 2005. They found that solar X-ray flares caused enhancements in the electron density by a factor of 5–6 during disturbed condition. The observed and estimated electron density profiles for quiet and disturbed conditions are represented in Fig. 33 for solar X-ray flare of 31 May 2003. It should be noted that the measured electron density profile cannot be reproduced completely by this model. This is due to the fact that while E and F layers in the Martian ionosphere at ∼90–110 km and ∼130–140 km are produced due to absorption of both solar X-ray and EUV radiations respectively, in this calculation only X-ray fluxes were used as input, which produce the E region of the Martian ionosphere. Haider et al. (2012) have reported that the modelled E layer peak height compares well with that of the MGS observations during quiet condition. MGS did not measure electron density during the maximum phase of the flare.
Fig. 33

Model calculation of electron densities at quiet (red line with star) and disturbed conditions (black line with circle) for SZA 78 on 31 May 2003. Eight profiles were observed by MGS on 31 May 2003. These profiles are averaged and plotted with error bars (blue color circle) in this figure for comparison with model calculations (from Haider et al. 2012)

5 Lower Atmosphere of Mars

The lower atmosphere of Mars is characterized by strong coupling between pressure, temperature, neutral density and winds. UV radiation up to 1950 Å can reach the surface of Mars. The cosmic rays are a major source of ionization in the lower atmosphere of Mars (cf. Haider et al. 2011). The lowest height (1–2 km) of the atmosphere is called the boundary layer where heat, momentum and mass are exchanged between the surface and the atmosphere. There is a strong temperature inversion in the nighttime at Mars due to radiative heat loss and the layer is shallow (Hess et al. 1976a, 1976b). The boundary layer height increases during the daytime, because of surface warming. The composition of the lower atmosphere is dominated by CO2 with other minor gases like N2, Ar, O2, H2, CO, H2O, O, O3, NO, NO2 and HNO3 (Molina-Cuberos et al. 2002; Haider et al. 2007, 2009c, 2011). The dust is an important constituent in the lower atmosphere of Mars and is always present in the background. It absorbs heat and changes the temperature structure of the atmosphere. The temperature and pressure have been measured in the lower atmosphere through remote sensing by Mariners 4, 6, and 7 and Mars 6. The Viking Landers provided the first in situ measurements of the Martian atmosphere. Recently the lower atmosphere of Mars has been monitored by thermal infrared spectrometer and radio occultation experiment onboard MGS (Smith et al. 2001; Hinson et al. 2004) and MEX (Grassi et al. 2005; Pätzold et al. 2005). Results of these measurements are given in the following sections.

5.1 Temperature and Pressure Measurements

The surface pressure and temperature of Mars varies with seasons (Kliore et al. 1965; Belton and Hunten 1966; Rasool et al. 1970; Seiff and Kirk 1977) due to condensation and sublimation of the atmospheric CO2 in the polar caps. The seasonal variations of atmospheric temperature and pressure were measured in situ by Viking 1 and 2 at landing sites (Hess et al. 1976a, 1976b). Figure 34 shows seasonal variability of the pressure observed by Viking 1 and 2 Landers on the surface of Mars. The first minimum of the pressure near solar longitude, Ls=150 occurs during southern winter when a great part of the atmosphere is trapped in the south polar cap. The second minimum, near Ls=430 corresponds to the northern winter much shorter and less cold than the southern winter (Mayo et al. 1977; Hourdin et al. 1993). Between northern autumn and northern winter seasons, the pressures are high because the eccentricity of Mars’ orbit leads to more intense evaporation in the summer at the southern pole. Atmospheric structure profiles have been observed on Mars Pathfinder during its descent in the early morning (3:00 am) on July 4, 1997 (Crisp et al. 2003). Pathfinder landed on Mars within 850 km of the Viking Landers in the late northern summer. Below 60 km altitude, the temperatures measured by Pathfinder are quite close to those measured by Viking Landers (Magalhaes et al. 1999). Vertical profiles of atmospheric density, pressure and temperature have also been observed in situ from the two Mars’ rovers, Spirit and Opportunity, which landed in the afternoon at local times ∼14:16 h and 13:13 h on 4 January and 25 January 2004 at coordinates −14.5N, 175.5E and −1.9N, 354.5E respectively (Withers and Smith 2006).
Fig. 34

Time evolution of the surface pressure observed by Viking 1/2 Landers during first 2587 three Martian years of the mission (Hess et al. 1976a, 1976b)

Using a remote sensing method MGS and MEX observed temperature, pressure and total density from radio occultation experiments in the northern and southern troposphere of Mars (Hinson et al. 1999; Pätzold et al. 2005). These and earlier observations have shown that up to 45 km the temperatures are highly variable and this variability depends on latitude, season and dust content of the atmosphere (Kliore et al. 1972; Smith et al. 2003; Haider et al. 2008; Tellmann et al. 2013). The Thermal Emission Spectrometer (TES) onboard MGS (cf. Smith 2004), Thermal Emission Imaging System (THEMIS) onboard Mars Odyssey (Smith 2009), Planetary Fourier Spectrometer (PFS) onboard MEX (Grassi et al. 2005; Tellmann et al. 2013) and Mars Climate Sounder (MCS) onboard Mars Reconnaissance Orbiter (MRO) (Hinson and Wilson 2004; Heavens et al. 2011) have also provided temperature profiles and their variability in the atmosphere of Mars. Withers and Smith (2006) have compared temperature-pressure profiles obtained from Viking 1/2 Landers, Mars Pathfinder and Mars rovers Spirit and Opportunity. They have found that Spirit’s temperature data shows large amplitude, long wavelength oscillations around 1 Pa. Pathfinder’s temperatures are relatively cold around 0.1–1 Pa. Viking 1/2 temperatures are relatively cold around 10 Pa, whereas Spirit’s temperatures are relatively warm around 10–100 Pa. Withers and Smith have suggested that these changes in the temperature profiles may occur due to changes in season, latitude, time of the day and dust content. Figure 35 represents a time series of temperature with solar longitude observed by TES and THEMIS experiments from MY24, Ls=104 to MY25, Ls=104 (Smith 2009). There is a good agreement between the two observations. The local time of THEMIS observation varied from 2:00 pm to 6:00 pm, while TES measured temperature daily at 2:00 pm. The atmospheric temperatures are periodic due to seasonal variability.
Fig. 35

Comparison of temperature variability measured by THEMIS and TES experiments onboard MEX and MGS respectively at different solar longitude (Ls). THEMIS experiment has measured this temperature at band 10. TES experiment provides a good estimate when there are no large dust storms (from Smith 2009)

5.2 Neutral Density and Composition

Using the ground-based experiment Kuiper (1952) inferred CO2 abundance ∼4.4×105 Dobson in the Martian atmosphere. After several years, Spinrad et al. (1963) and Kaplan et al. (1969) discovered water vapour and CO abundances ∼10 μm and ∼2.5 μm respectively. Ozone was discovered by Barth and Hord (1971) from UV spectrometer onboard Mariner 7 spacecraft. An upper limit of ∼5 % of N2 was reported from Mariner 6 and 7 measurements by Dalgarno and McElroy (1970). Mars 6 measurements indicated that 35 % of the Martian atmosphere might be a noble gas (Istomin and Grechnev 1976). Later Viking Landers measured neutral densities of O2, CO2, CO, N2, NO, Ar, Kr, Xe, and Ne in the lower atmosphere with neutral mass spectrometer (Nier and McElroy 1977; Biemann et al. 1976; Owen et al. 1976; McElroy et al. 1976; Owen et al. 1977). These measurements provided the first reliable composition of the Martian atmosphere, which is 95.3 % CO2, 2.7 % N2, 1.6 % Ar, 0.15 % O and 0.03 % H2O. Krasnopolsky et al. (1994) detected He in the lower atmosphere with mixing ratio of 1.1±0.4 ppm through airglow emission at 584 Å. Ar was measured with Gamma ray spectrometer onboard Mars Odyssey. This measurement showed that Ar is more at southern pole in comparison to that at northern pole (Sprague et al. 2012). The density of atomic hydrogen was measured from Lyman alpha airglow experiment onboard Mariner 6, 7 and 9 (Anderson 1974; Kong and McElroy 1977).

MGS and MEX have observed atmospheric density profiles in the troposphere of Mars with radio occultation experiment (cf. Hinson et al. 1999; Pätzold et al. 2005; Haider et al. 2011). Haider et al. (2008) constructed a model atmosphere of CO2, N2, Ar, O2, CO, H2, H2O, O, O3, NO, NO2, and HNO3 by multiplying mixing ratios of these gases with the total density observed by MGS. They used mixing ratios from Rodrigo et al. (1990) and Nair et al. (1994), who modelled the altitude profiles of neutral densities in the lower atmosphere of Mars. Figure 36 represents altitude profiles of neutral densities of 12 gases. This model atmosphere has been used by Haider et al. (2009b, 2010a, 2011) and Sheel et al. (2012) in their calculation of production rate, ion and electron densities of Mars. It should be noted that variations in mixing ratios can vary neutral density of minor constituents. The concentrations of major gases, however, are not much affected by any change in the mixing ratios. Haider et al. (2008) found that the neutral densities were larger in summer than that of winter by a factor of ∼2.
Fig. 36

Neutral atmospheric model of Sheel et al. (2012), which is used in the lower atmosphere of Mars. The red line shows the idealized atmosphere introduced by Sheel et al. (2012) for generalized simulation in their work

5.3 Dynamics of the Lower Atmosphere

Mars is a windy planet. Using radio occultation data of MGS, Haider et al. (2009c) have reported zonal wind perturbations in temperature and density of the troposphere of Mars during winter season at latitude range 60S–64S and during summer at latitude range 64N–67.3N. Figures 37 and 38 represent the nighttime zonal distribution of observed density at an altitude of 15 km during summer and winter seasons respectively. Solid lines represent model fit to the densities. The dotted and dashed lines are for lower and upper 0.95 confidence limits, respectively. The measurement uncertainties are also shown in the figures. A structure with two significant peaks in the density at ∼150E–175E and at ∼325E, can be seen in the northern troposphere during the summer season. Haider et al. have observed three prominent peaks in the density of southern troposphere at ∼50E, 180E, and 300E during the winter season. These peaks suggest that the Martian troposphere can be represented by wave number 2 in the northern high latitudes and wave number 3 in the southern high latitudes. It is clear from above that the tropospheric winds are different in the summer and winter seasons at high northern and southern latitudes. Seth et al. (2006a, 2006b) have argued that the Martian troposphere is changing due to forced and free oscillations. The forced oscillations consist of thermal tides and stationary waves, while free oscillations include atmospheric instabilities. The thermal tides consist of different mode numbers, associated with migrating and non-migrating tides, and have strongest influence in the troposphere by modulating the atmospheric density (Hollingsworth and Barnes 1996; Banfield et al. 2003).
Fig. 37

Zonal structure of density measured by MGS during summer at 15 km between longitude 0 and 360E and latitudes 64.7N and 67.3N (from Haider et al. 2009c)

Fig. 38

Zonal structure of density measured by MGS during winter at 15 km between longitude 0 and 360E and latitudes 60S and 64S (from Haider et al. 2009c)

5.4 Models of the Lower Atmosphere

Belton and Hunten (1969) used the photochemical model to calculate O2, O3, CO2 and CO in the lower atmosphere of Mars. The CO and O are produced in the atmosphere due to photolysis of CO2 at wavelength ≤ 2275 Å (McElroy and Donahue 1972). Parkinson and Hunten (1972) proposed an aeronomical model, which predicted a large amount of H2O2 whose photolysis could provide a source of OH to oxidize CO, leading to low abundances of O2 and CO on Mars and the corresponding stability of CO2. This was the first model which proposed the stability of Martian atmosphere. Later Krasnopolsky and Parshev (1979) developed a photochemical model in the lower atmosphere and obtained ozone ∼2.5×109 cm−3 at altitude ∼35 km. This model agreed well with the measurements made by UV photometer on Mars 5 spacecraft.

Using a 1-D model, the abundance of ozone as a function of height was estimated by Rodrigo et al. (1990). Their results were found in good agreement with the Mars 5 measurements. Later Krasnopolsky (2003a, 2003b) argued that the 1-D model has its own limitation since it cannot represent the latitudinal variations in the atmospheric densities. The 2-D model provided a better description of the interactions between dynamics, radiation and chemistry. Lefevre et al. (2004, 2008) used 3-D continuity equations to model the abundance of ozone on Mars. They have compared their model results with the satellite observations made by SPICAM experiment onboard MEX (Perrier et al. 2006) and Earth-based measurements carried by Hubble Space Telescope (Clancy et al. 1999) and the NASA Infrared Telescope (Fast et al. 2006). This study revealed that the model calculations using heterogeneous chemistry on the ice cloud is more important for quantitative agreement with the ozone levels observed by SPICAM (Lefevre et al. 2008). The 3-D models have a number of advantages over 1-D and 2-D models to represent the chemistry. First, the atmospheric transport is fundamentally three-dimensional and only 3-D models can properly calculate its effect on the distribution of chemical species. Second, the water vapour in General Circulation Model (GCM) can be represented at all spatial and time scales by the 3-D models.

6 Mars’ Lower Ionospheric Measurements

While relatively a large number of measurements have been made in the upper ionosphere of Mars, not much attention has been paid to its lower ionosphere. The lower ionosphere on Earth is generally inferred from the attenuation of HF waves. At Mars ionospheric reflections of radio waves are expected only for medium waves or VLF waves, since the plasma frequency is less than 104 cm−3. Such waves have a larger wavelength than the characteristic scale of the vertical gradient of electron density. Therefore the ray theory of wave propagation is not applicable. Consequently one needs to employ full wave treatment, which makes the estimation of electron density difficult in the lower ionosphere. The D layer at Mars has not been measured by in situ experiments. The existence of this layer has been found only in two electron density profiles measured by the radio occultation experiment onboard the Mars 4 and 5 orbiters (Savich and Samovol 1976). The progress in rocket observations and advanced technique of ground based sounding at Mars can bring the much desired information on the electrons and ions in the D region ionosphere. It is to be noted that in the upper ionosphere, electron density is equal to the total positive ion density, whereas in the lower ionosphere the electron density is less than the positive ion density. It is so because in the lower ionosphere, negative ions are more abundant than the electrons.

The presence of metallic plasma layers at about 80 km in the night time ionosphere of Mars was first indicated from the observations made by Mars 4 and Mars 5 (Savich and Samovol 1976). Later MGS and MEX confirmed the existence of sporadic layers in the daytime ionosphere between altitude 65 km and 100 km (Pätzold et al. 2005; Withers and Mendillo 2005; Withers et al. 2008). The ablation of meteoroids deposits metals in the atmosphere. These metals react with the ions and solar photons and produce sporadic low-lying plasma layers in the Martian ionosphere. In the fall of 2005, a dedicated meteor observing campaign was carried out by Mars Exploration Rover (MER) missions. To study the Martian surface geology the MER carried two rovers: Spirit and Opportunity. Domokos et al. (2007) analyzed Spirit data of a Panoramic Camera and estimated upper limit of meteoroid flux ∼4.4×10−6 meteoroids km−2 h−1 for mass larger than 4 g. Opportunity has found several meteorites on the surface of Mars ( We will discuss lower ionospheric measurements in detail in the following sections.

6.1 Mars 4 and Mars 5 Measurements

The dual frequency radio sounding of the Martian atmosphere above the dark surface of the planet was carried out during exits from behind the planet of spacecrafts Mars 4 and Mars 5 on 10 and 18 February 1974, respectively (Savich and Samovol 1976). Mars 4 occulted the ionosphere during the autumn season at solar zenith ∼127. The location was 90S, 236W and the local time was 03:30 h. Mars 5 occulted the ionosphere during spring season at solar zenith angle ∼106. The location was 38N, 214W and the local time was 04:30 h. Both measurements were carried out during a low solar activity period. Two electron density profiles were obtained from these measurements, which are shown in Fig. 39. Each measured profile shows the existence of two plasma layers, one at height ∼80 km and the other with a small shoulder of electron density ∼(1–2)×103 cm−3 at ∼25 km (Molina-Cuberos et al. 2008; Haider et al. 2009b). There can be a large error in these lower ionospheric measurements because the presence of plasma in the lower ionosphere is derived on the basis of formal inversion of the set of observations in symmetrical approximation. During the occultation measurements, the possible displacement of electron density due to this error is estimated to be ± 500 cm−3 (Savich and Samovol 1976). The second lowest peak is not very prominent and other published electron density profiles do not show this feature (Zhang et al. 1990). Pesnell and Grebowsky (2000) argued that the low altitude ionospheric structures observed by Mars 4/5 requires modelling and cannot be attributed to meteoroids. Withers et al. (2008; 2012) have also reported that these sporadic layers have not been detected in any of the other available electron density profiles.
Fig. 39

Electron densities measured by Mars 4 and Mars 5 in the nighttime ionosphere of Mars (from Savich and Samovol 1976; Verigin et al. 1991)

It should be noted that Mars 4/5 profiles were the first which indicated the presence of the nighttime ionosphere at Mars. The maximum value of electron concentration in the upper ionosphere of Mars during Mars 4 occultation was observed to be ∼4.6×103 cm−3 and ∼5×103 cm−3 for Mars 5 at heights of 110 km and 130 km, respectively. It can be easily shown that the night time ionosphere of Mars cannot be considered as the decay of the daytime ionosphere. The decrease of electron concentration after sunset is characterized by the time constant τ=1/(αNo), where No is the value of electron concentration at the moment of “switching off” the source of ionization, and α=2.55×10−7 cm3 s−1 (Mul and McGowan 1979) is the effective recombination coefficient. If we take the minimum observed value No∼104 cm−3, then τ∼ 400 sec. It is obvious that with this value of τ the ionospheric plasma would completely disappear soon after sunset. Thus, to explain the existence of night time ionosphere over Mars, it is necessary to consider the presence of an additional source of ionization, not connected directly with solar radiation. The source of the night time upper ionosphere of Mars has been described in detail in Sect. 3.2.3. The peaks that are produced at height ∼80 km and ∼25 km are due to metallic ablation and cosmic ray ionization respectively (cf. Zhang et al. 1990; Pätzold et al. 2005; Haider et al. 2008; Withers et al. 2008; 2012). The source and physical processes for the formation of these peaks in the lower ionosphere will be described in Sect. 7.

6.2 MGS Measurements and Meteoric Layers

As already mentioned, the radio occultation experiment onboard MGS obtained 5600 electron density profiles during the period 24 December 1998 to 9 June 2005. Withers et al. (2008) found meteoric layers in 71 out of these 5600 electron density profiles. They have reported mean peak electron density ∼(1.33±0.5)×1010 m−3 in the meteoric layer at a mean altitude of 91.7±4.8 km. Their analysis suggests that the characteristics of these meteoric layers vary with season, solar zenith angle and latitude. Later Pandya and Haider (2012) analyzed 1500 MGS electron density profiles to study the physical characteristics of meteoric plasma layers over Mars during the period January-June 2005. They found that 65 electron density profiles out of these 1500 profiles were strongly perturbed with peak densities ∼(0.5–1.4)×1010 m−3 at altitudes between 80 km and 105 km, presumably due to ionization of meteoric atoms. Using these profiles they examined TEC in the lower ionosphere of Mars (Fig. 40) and found that maximum values of TEC occurred on 21 January and 23 May 2005, when comets 2007 PL42 and 4015 Wilson-Harrington intersected the orbit of Mars from close distances of 1.49 AU and 1.17 AU, respectively. The TEC values were increased by a factor of 5–7 on these days. Pandya and Haider (2012) associated this significant increase with the meteor showers that were produced when Mars crossed the dust stream left along the orbits of these comets (Recently a new comet C/2013 A1 has been discovered at Siding Spring Observatory using Uppsala Southern Schmidt Telescope ( This comet will pass from Mars at minimum distance ∼97000 km in October 2014. Considering the size of coma ∼105 km, it can be said that Mars will pass in October 2014 through the gaseous envelop of the comet C/2013A1. Mars has a very tenuous atmosphere. Thus, the surface of Mars should have intensive bombardment of meteoroids during the encounter of this comet with the atmosphere of Mars).
Fig. 40

The monthly distribution of normalized TEC during January-June 2005 in the lower ionosphere of Mars. The meteor showers occurred on 21 January and 23 May 2005. TEC is normalized to the minimum value of each respective month between January and June 2005 (from Pandya and Haider 2012)

The closest approaches of comets 2007 PL42 and 4015 Wilson-Harrington with Mars are shown in Figs. 41 and 42, respectively. Most probably these comets left debris and dust particles in the Martian atmosphere, which ablazed and produced broad peaks in the lower ionosphere on 21 January and 23 May 2005. These responses of meteoric showers are clearly seen in the ionospheric profiles of 21 January and 23 May 2005, respectively shown in Figs. 43a and 43b. The peak electron densities during meteor showers were observed to be ∼1.06×1010 m−3 and 1.85×1010 m−3 at 99 km and 102 km respectively (Withers et al. 2008; Pandya and Haider 2012). The two meteor showers mentioned above were detected on Mars at different locations and at different times. The meteor shower of 21 January 2005 was observed at SZA=74.3, latitude 77.7 and longitude 197.2 during summer season (Ls=147.4). The meteor shower of 23 May 2005 was detected in the autumn season at Ls=216.2, SZA=84.9, latitude=65.1 and longitude 20.2E. In Figs. 43a and 43b, E and F peaks are produced due to solar X-ray and EUV radiation at altitude range ∼ 100–112 km and 125–135 km, with peak concentrations of 2.4×104 cm−3 and 8.4×104 cm−3 respectively (Withers et al. 2008; Mendillo et al. 2011; Fox and Weber 2012). The meteoric ablations have produced third ionization peaks ∼(1–2)×104 cm−3 below E layer in the altitude range 80–105 km. The electron density of this peak, however, is lower than that of E and F peaks which are produced by solar X ray and EUV radiation, respectively.
Fig. 41

Comet 2007 PL42 crossing Mars orbit from a closest distance 1.49 AU on 21 January 2005. Simulated view of comet 2007 PL42 and positions of Sun, Earth, Mars, Mercury and Venus are shown in this figure (

Fig. 42

Comet 4015 Wilson-Harrington crossing Mars orbit from a closest distance 1.17 AU on 23 May 2005. Simulated view of comet 4015 Wilson-Harrington and positions of Sun, Earth, Mars, Mercury and Venus are shown in this figure (

Fig. 43

(a) Altitude profiles of electron density measured by radio occultation experiment onboard MGS on 21 January 2005 and (b) 23 May 2005 (from Pandya and Haider 2012)

6.3 MEX Measurements and Meteoric Layers

Since 2004, radio occultation experiments onboard MEX have reported 557 vertical electron density profiles in the daytime and night time ionosphere of Mars. Pätzold et al. (2005) observed meteoroid layers in 8 % electron density profiles (i.e. 10 of 120 electron density profiles) obtained from MEX during the daytime ionosphere. Figure 44 shows an electron density profile observed by MEX on 18 April 2004 in the late evening at local time 17:00 h with solar zenith angle 85. This profile is nearly identical to the earlier measurements of the ionosphere made by MGS (Hinson et al. 1999) and has three ionization peaks. As mentioned earlier, the first and second layers (top and middle) are produced due to absorption of solar EUV and X-ray radiations, respectively. The third layer (lowest) appears occasionally between altitude range 80 km and 100 km with peak electron density of ∼104 cm−3 (Pätzold et al. 2005). Its origin has been attributed to ablation of meteor and charge exchange of magnesium and iron (Molina-Cuberos et al. 2003; Pesnell and Grebowsky 2000).
Fig. 44

Electron density profile measured by the MEX radio occultation experiment in the daytime on 18 April 2004. The F layer as M1 is present at 150 km, E layer as M2 is visible as a shoulder at 120 km, and third layer M3 is present at 85 km. The solid curve is a fit of Chapman function to the F and E layers. The open circles show the difference between the data and the fit (from Pätzold et al. 2005)

Withers et al. (2012) analyzed MEX occultation data and defined terminator ionosphere at 90<SZA<105 and dark ionosphere at SZA>105. They have found that the profiles of the dark ionosphere are different from those of the terminator ionosphere. These authors have reported two electron density profiles from the night time ionospheres of 21 August 2005 and 25 September 2005 observed with meteoroid layers at 76N, 264E and 42.7N, 319E, respectively. Figure 45 shows the electron density profile obtained on 25 September 2005 at solar zenith angle 121. This profile also represents three ionization layers in the night time ionosphere at altitudes ∼160 km, 120 km and 80–100 km. The first peak had been seen earlier in the night time ionosphere by Mars 4/5, Viking 1/2 and MGS spacecrafts (Savich and Samovol 1976; Zhang et al. 1990). This peak is known to be produced due to precipitation of solar wind electrons from the dayside to the nightside ionosphere across the terminator (Verigin et al. 1991; Haider et al. 1992, 2002; Fox et al. 1993; Haider 1997; Lillis et al. 2011). The second and third peaks were not detected in earlier observations at night time possibly due to the large uncertainty in the measurement. Mean uncertainty in MEX measurements for electron density is 6.32×108 m−3 (Pätzold et al. 2005), which is smaller than MGS by an order of magnitude. The mean value of uncertainty in MGS measurements for electron density is 4.6×109 m−3 (Hinson 2007; Tyler et al. 2007; Withers et al. 2008).
Fig. 45

Electron density profile observed by radio occultation experiment onboard MEX on 25 September 2005 in the nighttime ionosphere of Mars (Withers et al. 2012)

The values of the first and the second peak are lower by an order of magnitude in the night time as compared to those observed in the daytime ionosphere. In the night time, the height of the first peak is higher than the daytime peak by about 8–10 km. This implies that EUV energy is deposited at lower altitudes in the daytime ionosphere of Mars while solar wind energy is deposited at higher altitude in the night time ionosphere. The second layer is created at nearly the same height in the daytime and the night time ionosphere, but its source of ionization is different. The third plasma layer again is formed in the daytime and night time ionosphere at nearly the same altitude ∼80–100 km. It should be noted that the solar energy incident at Mars has periods such as daily rotation and seasons of Mars, a 27-day rotation of sun, an 11-year cycle of solar activity, and episodic events associated with solar flares. Therefore the height and density of first and second peak can change with variations in solar conditions. The concentration of the third plasma layer can change with intensity of meteor showers.

6.4 Composition and Chemistry of Meteoric Ions

The meteoroids enter in the atmosphere of Mars with a velocity similar to the orbital velocity and interact with its species (Yueha et al. 2002). The ablation of meteoroids deposits metals in the atmosphere, which react with the ions and solar photons producing low-lying plasma layers in the Martian ionosphere. Adolfsson et al. (1996) estimated the flux of meteoroids on Mars to be ∼50 % of that on Earth. They suggested that high speed (≥30 km/s) meteors will produce similar magnitudes of luminosities in the atmosphere of Earth and Mars. The height of maximum meteor intensity on Mars was found to be 50–90 km. On Earth it is 70–110 km. Christou and Beurle (1999) and Christou (2010) estimated the minimum distance of all comets/asteroids passing through the orbit of Mars and Earth. They found a total of 297 asteroids and 51 comets, which could approach Mars to within 0.2 AU compared to 156 and 24 respectively for the case of the Earth. This suggests that more meteor showers are visible in the environment of Mars than that are observed on Earth. Treiman and Treiman (2000) and Yueha et al. (2002) estimated velocity distribution of periodic comets around Mars. They found that 50 comets approached within 0.1 AU of Mars and these could be a potential source of dust and meteor showers at Mars.

Molina-Cuberos et al. (2003) developed a model for the calculation of meteoric ion concentration in the daytime and night time ionosphere of Mars. They reported that meteoric layers were formed of Mg+ and Fe+ with a magnitude of the order of 104 cm−3 at noon and decreased by two orders of magnitude at night. Figure 46 shows a schematic diagram of Mg metal chemistry. The chemistry of Fe is quite similar to that of Mg but has a different rate constant. Initially Mg and Fe atoms are lost by oxidation with O2 and O3 to form metal oxides as given below:
$$\begin{aligned} \mathrm{Mg} + \mathrm{O}_{3} \rightarrow& \mathrm{MgO} + \mathrm{O}_{2} \\ \mathrm{Mg} + \mathrm{O}_{2} + \mathrm{M} \rightarrow& \mathrm{MgO}_{2} + \mathrm{M} \\ \mathrm{Fe} + \mathrm{O}_{3} \rightarrow& \mathrm{FeO} + \mathrm{O}_{2} \\ \mathrm{Fe} + \mathrm{O}_{2} + \mathrm{M} \rightarrow& \mathrm{FeO}_{2} + \mathrm{M} \end{aligned}$$
The oxidation of ozone with Mg and Fe is a minor process for the production of MgO and FeO. These metals also undergo three body reactions and produce large quantities of MgCO3 and FeCO3.
Fig. 46

Schematic diagram of Magnesium chemistry in the Martian ionosphere (from Molina-Cuberos et al. 2003)

Molina-Cuberos et al. (2003) showed that the destruction of Mg and Fe with atmospheric ion \(\mathrm{O}_{2}^{+}\) is more important than the photoionization of these neutral metals. This loss process produces Mg+ and Fe+ at about 80 km. Recently Haider et al. (2013) have estimated densities of 21 ions (viz. Mg+, Fe+, Si+, MgO+, \(\mathrm{MgCO}_{2}^{+}\), \(\mathrm{MgO}_{2}^{+}\), \(\mathrm{MgN}_{2}^{+}\), FeO+, \(\mathrm{FeO}_{2}^{+}\), \(\mathrm{FeN}_{2}^{+}\), \(\mathrm{FeCO}_{2}^{+}\), SiO+, \(\mathrm{SiCO}_{2}^{+}\), \(\mathrm{SiN}_{2}^{+}\), \(\mathrm{SiO}_{2}^{+}\), \(\mathrm{CO}_{2}^{+}\), \(\mathrm{N}_{2}^{+}\), O+, \(\mathrm{O}_{2}^{+}\), CO+ and NO+) in the night time ionosphere of Mars. Their model showed that the atmospheric ions \(\mathrm{CO}_{2}^{+}\), \(\mathrm{N}_{2}^{+}\), O+, CO+, \(\mathrm{O}_{2}^{+}\) and NO+ are produced above 100 km and are formed due to solar wind electron and proton impact ionization. The metallic ions are formed between 50 km and 100 km due to the ablation of micrometeoroids. Figure 47 represents density profile of meteoric ions in the night time ionosphere of Mars. Haider et al. have solved equations of motion, ablation and energy for the modeling of ion production rates due to impact of meteoroids as given below:
$$\begin{aligned} \operatorname{Cos} \theta \frac{dV}{dz} =& \frac{\varGamma A \rho V}{m^{1/3} \delta^{2/3}} \end{aligned}$$
$$\begin{aligned} \operatorname{Cos} \theta \frac{d m}{dz} =& - \frac{4 A m^{2/3} C_{1}}{\delta^{2/3} T^{1/2} V} e^{ - C_2 /T} - \frac{\varLambda_{s} A m^{2/3} \rho V^{2}}{2 Q \delta^{2/3}} \end{aligned}$$
$$\begin{aligned} \operatorname{Cos} \theta \frac{d T}{dz} =& \frac{4 A \rho V^{2}}{8 C\delta^{2/3} m^{1/3}} (\varLambda - \varLambda_{s}) - \frac{4 A \sigma T^{4}}{C \delta^{2/3} V m^{1/3}} - \frac{4 A C_{1} Q}{C \delta^{2/3} T^{1/2} m^{1/3} V} e^{ - C_2 /T} \end{aligned}$$
where V is the meteoroid velocity, A is the shape factor, δ is the meteoroid density, ρ is the atmospheric density, m is the mass of meteor substance, Γ is drag coefficient, Λ is the heat transfer coefficient, Λs is the sputtering coefficient, Q is the energy of evaporation, σ is the Stefan-Boltzman constant, θ is the entry angle and C is the heat capacity of meteoroid material. Parameters C1 and C2 define the dependence of the evaporation rate on temperature. The ion production rates so obtained were used later in the continuity equation (Eq. (1)) to calculate the density of meteoric metals and ions in the lower atmosphere of Mars.
Fig. 47

Calculated densities of metallic ions due to meteor ablation in the nighttime ionosphere of Mars (from Haider et al. 2013)

7 Modeling of the Lower Ionosphere of Mars

The first theoretical study of the lower ionosphere of Mars was carried out by Whitten et al. (1971). They considered ionization by cosmic rays and solar radiation on the dayside ionosphere of Mars. Later Molina-Cuberos et al. (2002) calculated electron density in the daytime lower ionosphere of Mars. Haider et al. (2007, 2009b, 2009c) modelled electron density in the daytime and night time ionosphere of Mars due to the absorption of solar wind electron, solar EUV and galactic cosmic rays between altitudes 0 and 220 km. They reported that daytime ionosphere can be divided into D, E, and F layers at altitude ranges ∼ 25–35 km, ∼ 100–112 km and ∼125–145 km owing to the impact of galactic cosmic rays, X-rays and solar EUV radiation respectively. Due to the absence of X-ray radiation, E layer is not produced in the night time ionosphere. The F and D layers are produced in the night time ionosphere at altitudes ∼130–140 km and ∼30 km due to absorption of solar wind electrons and galactic cosmic rays respectively. The lower ionosphere models have calculated production rates, densities of positive ions and negative ions (Molina-Cuberos et al. 2002; Haider et al. 2007, 2009b). These models couple ion-neutral, electron neutral, dissociation of positive and negative ions, electron detachment, ion-ion, and ion-electron recombination processes. In the chemistry of positive ions, it is found that hydrated hydronium ions, H3O+(H2O)n for n=1,2,3 and 4 dominate below 60 km, while NO+ and \(\mathrm{O}_{2}^{+}\) are major ions above this altitude (Haider et al. 2009b). In the chemistry of negative ions, one finds that water cluster \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\) and \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\) for n=1 and 2 are dominant ions below 40 km (Haider et al. 2009b). Above this altitude electrons play an important role in the Martian ionosphere.

First modelling of the meteoric ionization in the lower ionosphere of Mars was carried out by Pesnell and Grebowsky (2000). Using the photochemical equilibrium model, they calculated the deposition rate and density of Mg+ due to ablation of meteoroids in the Martian atmosphere. Pesnell and Grebowsky (2000) calculated the density of Mg+ to be ∼104 cm−3 at about 80 km due to ablation of meteoroids in the Martian atmosphere. They found that the maximum concentration of the Mg+ layer was not very sensitive to the velocity of the incoming micrometeoroids. Later Molina-Cuberos et al. (2003, 2008) developed a model of the ionosphere of Mars between 60 and 120 km for the daytime and night time conditions. In this model, ion and neutral chemistry of Mg, Fe and Si was included for the calculation of densities of neutral and ionic metals. This model included various processes viz. ablation of meteoroids, metals diffusion in the atmosphere, photoionization and chemistry of the ion and neutrals including charge exchange. Recently Haider et al. (2013) have extended this model and estimated the densities of metallic and atmospheric ions in the night time ionosphere between 50 km and 220 km.

7.1 Cosmic Ray Induced Ion Production Rates

The high-energy cosmic rays propagate through the Martian atmosphere producing nucleonic cascades. The impact of primary cosmic rays onto the atmospheric gases produces protons, neutrons and pions. The neutral pions quickly decay to gamma rays and their contribution to the energy deposition is very important in the lower part of the atmosphere. Near the mesosphere, the maximum ion production rates are controlled by protons. The charged pions decay to muons, which do not decay, and their energy is transferred to the surface upon reaching the ground. The flux of incident cosmic rays has been calculated to be ∼103 to 10−5 particles m−2 s−1 GeV−1 Ster−1 at energy interval 1 to 1000 GeV (O’Brien et al. 1996). Using the energy loss method, Haider et al. (2007, 2008, 2009b) calculated production rates of ions \(\mathrm{CO}_{2}^{+}\), \(\mathrm{N}_{2}^{+}\), Ar+, \(\mathrm{O}_{2}^{+}\), CO+, O+, N+, \(\mathrm{O}_{3}^{+}\), NO+, \(\mathrm{NO}_{2}^{+}\), \(\mathrm{HNO}_{3}^{+}\) and H2O+ due to cosmic ray ionizations using the equation:
$$ \biggl( \frac{dE}{dh} \biggr)_{\mathit{ion}} = 4\pi r_{0}^{2} \frac{m_{0}c^{2}}{\beta^{2}}NZ \biggl\{ \ln \beta \biggl( \frac{E + m_{0}c^{2}}{I} \biggr) \biggl( \frac{E}{m_{0}c^{2}} \biggr)^{\frac{1}{2}} - \frac{1}{2} \beta^{2} \biggr\} $$
where E is the energy, I is ionization potential, N is the neutral density r0 is classical electron radius with \(4\pi r_{0}^{2} = 1.0 \times 10^{ - 24}\) cm2/electron; m0c2=0.51 MeV, Z is the atomic number, β2=(V/c)2=1−[(E/m0c2)+1]−2 and c is the velocity of light. Using Eq. (8) the ion production rate at height h and solar zenith angle χ can be given by:
$$ p(h,\chi ) = \frac{1}{Q}\int_{E} \int _{\varOmega} (dE/dh) F(\chi,E,\varOmega )d\varOmega dE $$
where Q=35 eV, is the energy required for the formation of an electron ion pair, F is the total differential flux of the galactic cosmic rays and Ω is the spatial angle. In this calculation neutral model atmosphere consisting of CO2, N2, N, Ar, O2, CO, H2O, O3, O, NO, NO2 and HNO3, as mentioned earlier, was taken from various sources (cf. Rodrigo et al. 1990; Nair et al. 1994; Owen et al. 1977; Haider et al. 2009b).

7.2 Chemistry of D Region Ionosphere

The D region in the lower ionosphere of Earth is formed due to following ionization sources: (1) solar Lyman α (1216 Å) ionizing the minor constituent NO, (2) solar X-ray (λ<8 Å) ionizing N2 and O2, (3) cosmic ray ionizing all atmospheric constituents and (4) photoionization of metastable O2( 1Δg) by solar EUV radiation (λ<1118 Å) (e.g. Kelley 1989; Hargreaves 1992; Brasseur and Solomon 2005). The densities of ten positive ions (viz. \(\mathrm{O}_{2}^{+}\), \(\mathrm{CO}_{2}^{+}\), \(\mathrm{O}_{2}^{+}\mathrm{CO}_{2}\), H3O+, H3O+H2O, H3O+(H2O)2, \(\mathrm{CO}_{2}^{+}\mathrm{CO}_{2}\), \(\mathrm{O}_{2}^{+}(\mathrm{CO}_{2})_{2}\), H3O+(H2O)3, H3O+(H2O)4) and total ion density are shown in Fig. 48 (Sheel and Haider 2012). In the chemistry of positive ions, hydrated hydronium ions, H3O+(H2O)n for n=1,2,3 and 4 dominate below 70 km and \(\mathrm{O}_{2}^{+}\) ions dominate above this altitude. Molina-Cuberos et al. (2002) developed an ion-neutral model for the study of the positive and negative ion densities in the lower ionosphere. They found that three body reactions are important below 70 km, while charge exchange dominates above this altitude. Later this model was extended by Haider et al. (2007, 2008, 2009b) and Sheel et al. (2012) to study the chemistry of the upper and lower ionosphere of Mars simultaneously. In the latest model, \(\mathrm{O}_{2}^{+}\) and \(\mathrm{CO}_{2}^{+}\) ions are produced initially because of the impact of galactic cosmic rays between 0 km and 80 km (Sheel and Haider 2012). Later nearly 100 % \(\mathrm{O}_{2}^{+}\mathrm{CO}_{2}\) is formed by three body reactions, which is then fully destroyed by water vapour in the formation of H3O+. The ion H3O+ is lost further by three body reactions in the formation of hydrated ions H3O+(H2O)n for n=1–4.
Fig. 48

Altitude profiles of densities of positive ions. The solid line represents the total density of all positive ions (from Sheel and Haider 2012)

The estimated densities of ten negative ions (viz. \(\mathrm{O}_{2}^{-}\), \(\mathrm{CO}_{3}^{-}\), \(\mathrm{CO}_{4}^{-}\), \(\mathrm{NO}_{2}^{-}\), \(\mathrm{CO}_{3}^{-}\mathrm{H}_{2}\mathrm{O}\), \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\), \(\mathrm{NO}_{2}^{-}\mathrm{H}_{2}\mathrm{O}\), \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\), \(\mathrm{NO}_{3}^{-}\mathrm{H}_{2}\mathrm{O}\), \(\mathrm{NO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\)) electron and total negative ions are shown in Fig. 49 (Sheel and Haider 2012). In the chemistry of negative ions, water clusters of \(\mathrm{CO}_{3}^{-}\) and \(\mathrm{NO}_{2}^{-}\) (i.e. \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\), \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\) for n=1 and 2) are major ions below 40 km. Above this altitude electrons play an important role in the lower ionosphere. The maximum electron density occurs at ∼30 km due to the high efficiency of electron attachment to Ox molecules. This peak has been identified as D layer in the lower ionosphere of Mars (Whitten et al. 1971; Molina-Cuberos et al. 2002; Haider et al. 2007, 2008, 2009b). In the D region, the negative ions O and \(\mathrm{O}_{2}^{-}\) are produced by electron capture and molecular oxygen respectively (Haider et al. 2008). Later ion \(\mathrm{CO}_{4}^{-}\) is destroyed by atomic oxygen and \(\mathrm{CO}_{3}^{-}\) ion is produced. The loss of \(\mathrm{CO}_{3}^{-}\) with H2O has been found to be a major source of \(\mathrm{CO}_{3}^{-}\mathrm{H}_{2}\mathrm{O}\) and \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\). These reactions produce 100 % \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\) at all heights in the troposphere of Mars (see Fig. 49). The ion \(\mathrm{CO}_{3}^{-}\) is also destroyed by NO and NO2 forming \(\mathrm{NO}_{2}^{-}\) and \(\mathrm{NO}_{3}^{-}\) respectively. Later \(\mathrm{NO}_{2}^{-}\) and \(\mathrm{NO}_{3}^{-}\) are hydrated by three body reaction with water vapour and produce second important ions \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\) and \(\mathrm{NO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\), which are dissociated by neutral collisions.
Fig. 49

Altitude profiles of densities of negative ions. The dashed line represents the total density of all negative ions. The electron density is shown by solid line (from Sheel and Haider 2012)

7.3 Zonal Variability in the D Region Ionosphere

Mars has dramatic surface topography, huge diurnal variations in surface solar heating and temperatures, and strong surface winds. These factors act together to generate large amplitude waves occurring on a wide range of spatial and temporal scales. The diurnal variations in surface heating, coupled with strong longitudinal variations in surface topography, excite enormous tidal structures termed as non-migrating tides. These tides are linked both to variable terrain and to solar forcing. They have very large amplitudes and thus penetrate into the Martian lower atmosphere. In addition, Mars is a very windy planet, and the wind is composed of gravity waves, tidal waves and stationary waves (Imamura and Ogawa 1995; Bougher et al. 2001; Haider et al. 2006; Cavalie et al. 2008). The propagating winds can lead to variations in tropospheric temperature and density. MGS has observed temperature, density and pressure from radio occultation experiment for over three Martian years (Hinson et al. 2004). Haider et al. (2009b) have reported zonal wave structures in the night time tropospheric density and temperature and in the D region ionosphere of Mars. They calculated production rates, ion and electron densities between longitudes 0 and 360 due to impact of galactic cosmic rays on the night time troposphere. Figure 50 represents the calculated peak electron densities for the latitude range 64.7N–67.3N for summer season as a function of longitude. In this figure there are two prominent zonal peaks at 110E and 240E, which indicate a longitudinal delay of about 60E in the calculated electron density values and the neutral density observations (see Fig. 37). In the following we shall explain this delay as proposed by Haider et al. (2009b).
Fig. 50

Longitudinal distribution of D peak electron densities for summer season at latitude range 64.7N–67.3N in the night time ionosphere of Mars. The solid line represents the best fit to the calculation. The dotted and dashed lines represent lower and upper 0.95 confidence limits (from Haider et al. 2009c)

In the D region ionosphere, there are a significant number of positive and negative ions. Under chemical equilibrium conditions, electrical neutrality is required as Ne+N=N+, where Ne, N and N+ are the concentrations of electrons, negative ions and positive ions respectively. Since the negative and positive ions may also combine with each other, the overall balance between production and loss is now expressed as q=αeNeN++αiNN+, where q is production rate, αe and αi are the recombination coefficients for the reactions of positive ions with electrons and negative ions, respectively. If the ratio between negative ions and electron concentrations is represented by λ, then in terms of λ, N=λNe and N+=(1+λ)Ne and thus \({q} = (1+\lambda)(\alpha_{\mathrm{e}} + \lambda \alpha_{\mathrm{i}}){N}_{\mathrm{e}}^{2}\), which becomes as \({q} = (1+\lambda)\alpha_{\mathrm{e}}{N}_{\mathrm{e}}^{2}\) for λαiαe. Thus, the equilibrium electron density in the D region ionosphere is proportional to the square root of the ion production rate and its magnitude is changed to \({q}=\alpha\mathrm{e} {N}_{\mathrm{e}}^{2}\) when λ≪1. This equation is applicable in the upper ionosphere, where negative ions are absent. Haider et al. (2009b) argued that the longitudinal delay mentioned above is associated with the value of λ, whose contribution is about 10 % in the peak electron density of D region ionosphere. With this value of λ, the position of the crests and troughs in the peak electron density will be shifted with respect to that obtained in the longitudinal distribution of neutral density and thus production rates.

7.4 Ion-Dust Aerosol Model and Atmospheric Conductivity

In the absence of momentum term Eq. (9) can be represented for positive and negative ion concentrations of atmospheric gases and charged aerosol as:
$$\begin{aligned} \displaystyle\frac{d\rho_{i}^{ +}}{dt} = q_{i}^{ +} - \sum _{k} K^{i} \rho_{k} \rho_{i}^{ +} - \alpha_{ii}\rho_{i}^{ +} \rho_{i}^{ -} - \alpha_{e}\rho_{i}^{ +} \rho_{e} - \varGamma^{i}\rho_{i}^{ +} - \alpha_{1}A_{i}^{ -} \rho_{i}^{ +} - q_{A}^{ +}& \end{aligned}$$
$$\begin{aligned} \displaystyle\frac{d\rho_{i}^{ -}}{dt} = q_{i}^{ -} - \sum _{k} K^{i} \rho_{k} \rho_{i}^{ -} - \alpha_{ii}\rho_{i}^{ +} \rho_{i}^{ -} - \varGamma^{i} \rho_{i}^{ -} - \alpha_{2}A_{i}^{ +} \rho_{i}^{ -} - q_{A}^{ -}& \end{aligned}$$
$$\begin{aligned} \displaystyle\frac{dA_{i}^{ +}}{d t} = q_{A}^{ +} - \beta_{e}A_{i}^{ +} \rho_{e} - \alpha A_{i}^{ +} A_{i}^{ -} - \alpha_{2} A_{i}^{ +} \rho_{i}^{ -}& \end{aligned}$$
$$\begin{aligned} \displaystyle\frac{dA_{i}^{ -}}{dt} = q_{A}^{ -} - \alpha A_{i}^{ +} A_{i}^{ -} - \alpha_{1}A_{i}^{ -} \rho_{i}^{ +}& \end{aligned}$$
where \(\rho_{i}^{ +}\) and \(\rho_{i}^{ -}\) are the positive and negative ion concentrations, respectively, in the presence of aerosols, \(A_{i}^{ +}\) and \(A_{i}^{ -}\) are the positive and negative charged aerosol concentrations, respectively, βe and αe are the charged aerosol and ion-electron recombination coefficients, respectively, ρe is the electron concentration, \(q_{i}^{ +}\) and \(q_{i}^{ -}\) are the total ion production rates of positive and negative ions, respectively, \(q_{A}^{ +}\) and \(q_{A}^{ -}\) are the production rates of positive and negative charged aerosols, respectively, α is the charged aerosol-charged aerosol recombination coefficient, α1 is the positive ion and negative-charged aerosol recombination coefficient, α2 is the negative ion and positive charged aerosol recombination coefficient, Γi is the coefficient rate for the loss of ith ion by photon collision, αii is the ion-ion recombination coefficient, Ki is the rate coefficient of the reaction to remove the ion species i by the reaction with neutral concentration ρk of species k.

Equations (10)–(13) couple ion-neutral collisions, electron-neutral collisions, dissociation of positive and negative ions, electron detachment of anions, ion-dust attachment, ion-ion, and ion-electron recombination processes. This model by Haider et al. (2010b) consists of twelve gases (viz. CO2, N2, N, Ar, O2, CO, H2O, O3, O, NO, NO2 and HNO3) and 140 chemical reactions taken from Molina-Cuberos et al. (2002) and Haider et al. (2007, 2008). The neutral aerosol concentrations have been computed from dust mixing ratio profiles taken from Montabone et al. (2006). Whitten et al. (2007) and Michael et al. (2007) have derived a steady state recurrence relation to compute the build-up of electric charge on aerosol particles due to collision with positive and negative ions and electrons. Haider et al. (2010b) used this recurrence relation and allowed to vary the charges on aerosols from −30 to +30.

The conductivity depends on the number density of charged particles according to the relation, \(\sigma = e ( \sum_{i} \mu_{i}^{ +} \rho_{i}^{ +} + \sum_{i} \mu_{i}^{ -} \rho_{i}^{ -} )\), where σ is the total ion conductivity, μ+ and μ are the positive and negative ion mobility, respectively, and e is the elementary charge. The mean ion mobility is calculated as 0.02 m2 V−1 s−1 in the Martian ionosphere (cf. Borucki et al. 1982; Michael et al. 2007). By using this value of the mobility and the concentrations Sheel and Haider (2012) have calculated the total ion conductivities in the presence and absence of dust aerosols in the troposphere of Mars as shown in Fig. 51. During dust storms, the total ion conductivity is reduced by ∼2 orders of magnitude near the surface of Mars, as we shall see in the next section.
Fig. 51

Altitude profiles of total ion conductivities in presence and absence of dust aerosols (from Haider et al. 2010b)

7.5 Effect of Dust Storms in the D Region Ionosphere

Recently, the THEMIS experiment onboard Mars Odyssey has observed the dust optical depth in the thermal infrared for a period of about four years (February 2002 to July 2009). These data are shown in Fig. 52 at intervals of 5 in Ls (areocentric longitude) and 5 in latitude. The optical depths are averaged over longitude. The local time of this observation varies from 16:00 to 18:00 h. A major dust storm occurred in the month of August 2007 during southern summer at Ls=265–280 (Smith 2009). The location of the dust storm is shown in Fig. 52 by a rectangular box. Haider et al. (2010b) have used dust opacity values of this dust storm to study the effect of dust aerosols in the D region of the Martian ionosphere. They solved Eqs. (10) to (13) for the calculation of positive and negative ions in the presence and absence of aerosols in the Martian troposphere. The number densities of 12 positive ions (viz. \(\mathrm{O}_{2}^{+}\), \(\mathrm{CO}_{2}^{+}\), \(\mathrm{CO}_{2}^{+}\mathrm{CO}_{2}\), \(\mathrm{O}_{2}^{+}\mathrm{CO}_{2}\), H3O+, H+(H2O)2, H+(H2O)3, H+(H2O)4, H+(H2O)5, H3O+OH, \(\mathrm{O}_{2}^{+}\mathrm{H}_{2}\mathrm{O}\) and \(\mathrm{O}_{2}^{+}(\mathrm{CO}_{2})_{2}\)) with dust and without dust aerosols are shown in Figs. 53b and 53a, respectively. The number densities of 13 negative ions (viz. \(\mathrm{O}_{2}^{-}\), \(\mathrm{CO}_{4}^{-}\), \(\mathrm{O}_{3}^{-}\), \(\mathrm{CO}_{3}^{-}\), O, \(\mathrm{NO}_{3}^{-}\), \(\mathrm{NO}_{2}^{-}\), \(\mathrm{CO}_{3}^{-}\mathrm{H}_{2}\mathrm{O}\), \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\), \(\mathrm{NO}_{2}^{-}\mathrm{H}_{2}\mathrm{O}\), \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\), \(\mathrm{NO}_{3}^{-}\mathrm{H}_{2}\mathrm{O}\) and \(\mathrm{NO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{2}\)) with dust and without dust aerosols are shown in Figs. 54b and 54a respectively. It can be seen in these figures that the concentration of water cluster ions H+(H2O)n, \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\) and \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{{n}}\) gets reduced by 2 orders of magnitude in the presence of dust aerosols. Haider et al. (2010b) have argued that the lower ionosphere of Mars can be perturbed significantly during the massive dust storms and a large hole in the ion concentration may appear during the lifetime of a dust storm.
Fig. 52

Seasonal variability of dust optical depths observed by THEMIS onboard Mars Odyssey. The data are from February 2002 to July 2009. The rectangular box shows the location of a major dust storm. The Martian Year (MY) is also shown at the top (from Smith 2009; Haider et al. 2010b)

Fig. 53

Altitude profiles of positive ion densities calculated by the model for two cases of positive ions: (a) without dust aerosol (b) with dust aerosol (from Haider et al. 2009a)

Fig. 54

Altitude profiles of negative ion densities calculated by the model for two cases of negative ions: (a) without dust aerosols and (b) with dust aerosol (from Haider et al. 2010b)

8 Discussion

There is a strong focus on atmospheric measurements on Mars, from early Viking descent soundings to the more recent aerobraking and limb sounding by MGS, Mars Odyssey, MEX and MCS measurements aboard MRO. Collectively, measurements by these missions have demonstrated the strong variability in the Martian atmosphere on short-term to seasonal time scales, small-scale to planetary spatial scales and from the Martian surface to the thermosphere/exosphere. The atmospheric features responsible for these variations are solar forcing, major dust storms and the atmospheric wave motions. As on Earth, the tidal and planetary waves have also a major influence on the lower part of the Martian atmosphere (Bougher et al. 2001; Seth et al. 2006a, 2006b; Haider et al. 2011). The gravity waves contribute to much of the variability in wind, temperature and density at altitudes starting from the Martian surface into the thermosphere/exosphere (Heavens et al. 2011). An instrument that overcomes the above liabilities by providing both sensitivity to small spatial scales and sensitivity to the relevant wave perturbations in temperature, extending down to the Martian surface and into the thermosphere, is needed on future missions. The key questions that can be addressed comprehensively by such an instrument include: (1) What are the seasonal and short-term structure and variability of Martian atmosphere? (2) What are the roles of dust storms and planetary boundary layer in the large-scale dynamics, structure and variability at higher altitudes?, and (3) What are the implications of large density and wind fluctuations near the surface, and how do they vary with seasons and years (long term)?

At present, measurements on the upper ionosphere of Mars are primarily limited to the radio occultation profiles obtained from the Mariner 4, 6, 7 and 9 (Barth et al. 1971), Mars 2, 4, 5 and 6 (Moroz 1976), Viking 1 and 2 (Nier and McElroy 1977), MGS (Hinson et al. 1999) and MEX (Pätzold et al. 2005). This experiment measured E and F layers at altitude range ∼100–112 km and ∼125–145 km (Mendillo et al. 2003; Mahajan et al. 2007; Fox and Yeager 2006; Withers 2009; Haider et al. 2009a, 2009b). Pesnell and Grebowsky (2000) and Molina-Cuberos et al. (2003) predicted meteoric layers at altitude range ∼85–95 km. This prediction was confirmed from radio occultation measurements onboard MGS and MEX (Pätzold et al. 2005; Withers et al. 2008). All of these measurements were carried out at mid-and high-latitudes (cf. Nagy et al. 2004; Haider et al. 2011). Mars’ ionosphere has not been observed at low latitude (≤30). The challenge for the future is to develop theories and experiments to study the dynamics of the low latitude ionosphere. The physics of the low latitude ionosphere could be very different from the middle-and high-latitude ionosphere as discovered from the satellite and ground-based observations in the case of Earth’s ionosphere. Therefore, it is necessary to look for opportunities for obtaining observations at low latitude regions.

We have very few measurements in the lower ionosphere (≤ 60 km) of Mars. However, there are several models, which have predicted D layer at altitude ∼25–35 km (Whitten et al. 1971; Molina-Cuberos et al. 2002; Haider et al. 2009c; 2010b). In the D region ionosphere there are a significant number of cluster ions, which are produced due to cosmic ray ionizations. In-situ measurements of the cluster ion densities in the D region can be performed from the future Mars Landers. Sounding rockets carrying a quadrupole ion mass spectrometer have been used to measure the densities of cluster ions in the Earth’s lower ionosphere (Kopp and Hermann 1984). This experiment can observe the density of cluster ions in the lower ionosphere of Mars. D layer in the Earth’s ionosphere is formed at altitude ∼60–85 km mainly owing to solar Lyman α radiation ionizing minor constituent NO. The ionization caused by solar Lyman α will be almost ineffective on Mars because the density of NO is lower by ∼2–3 orders of magnitude in the Martian mesosphere as compared to that observed in the mesosphere of Earth (cf. Aikin 1968; Hargreaves 1992; Haider et al. 2009c).

The theory of the photochemistry and ion chemistry of Mars is now well developed and tested for specific times and locations as observed by the spacecrafts. The thermal structure needs to be understood better. There is now the beginning of a thermal general circulation model which now has been tested at a few points. The challenge for the future is to develop theories and models to describe the diurnal and seasonal changes that occur over the entire Martian globe. The solar energy output varies with various periods such as the daily rotation of Mars; the seasons of Mars; the 27-day rotation period of the sun; the 11 year cycle of solar activity; and episodic events associated with solar flares. Tests of aeronomic models with the above mentioned changing conditions will be challenging, but the reward will be enormous. We should then have the ability to understand the past and to predict the future with greater confidence. By developing such models we can also understand the history of Mars’ atmosphere/ionosphere, climate, liquid water and planetary habitability.

9 Summary and Conclusion

In this paper we have reviewed experimental and modeling results obtained so far on the upper and lower atmosphere of Mars. We described in detail atmospheric and ionospheric structures and their (1) disturbances, (2) composition, (3) chemistry and (4) dynamics. The aim of this review was to provide an overview of our current understanding of various aspects of Mars’ atmosphere and ionosphere at a scientific level which is comprehensive enough for students and researchers working in planetary aeronomy. Future missions like ExoMars by ESA, MAVEN by NASA, Mangalyaan by ISRO and MELOS by JAXA are in pipeline to explore Mars. MAVEN and Mangalyaan will fly in November 2013. MAVEN will explore the planet’s upper atmosphere, ionosphere and their interaction with solar wind. The objective of the Mangalyaan mission is to study methane, D/H ratios, dust storms, surface features and atmospheric composition in the upper atmosphere/exosphere of Mars. ExoMars will fly two missions, the first in 2016 and the second in 2018. The first mission will carry Trace Gas Orbiter to search for evidence of methane and other atmospheric gases that could provide signatures of biological activity. The 2018 mission will land a rover on Mars with scientific instruments to investigate the Martian environment. The MELOS mission would carry an orbiter and four Landers. The orbiter would study the atmosphere, its interaction with solar wind and the atmospheric weather. The four Landers will study surface, astrobiology, interior and sample returns. The MELOS is scheduled to fly in July 2020. It is expected that new missions will help us answer several key questions concerning the upper and lower atmosphere/ionosphere reviewed in this paper. The major questions that are to be understood from these missions are three-fold: (1) what are current escape rates of neutral gases and ions to space? It is believed that over millions of years, Mars’ core cooled and its magnetic field decayed, allowing the solar wind to sweep away ∼99 % of its atmosphere, (2) the existence of life on Mars has long been questioned. Recent detection of methane in the Martian atmosphere has generated much interest among the scientific community as it suggests the possibility of biological activities. But the presence of methane in the Martian atmosphere can also be explained in terms of geological activities. To find out the origin of methane, whether it is biotic or abiotic, may solve some of the most important questions regarding the existence of life forms on Mars, and (3) what is the ratio of D/H in the Martian atmosphere? The determination of D/H ratio allows us to understand the history of Mars’ atmosphere, climate, liquid water and planetary habitability.



We acknowledge C.T. Russell for his encouragement to write this review paper. One of the authors (S.A. Haider) is thankful to Bhavin Pandya for providing graphics support during the preparation of this manuscript. K.K. Mahajan is thankful to the Indian National Science Academy for the award of INSA Honorary Scientist scheme.


  1. M.H. Acuña et al., Magnetic field and plasma observations at Mars: Initial results of the Mars Global Surveyor Mission. Science 279, 1676–1680 (1998) ADSGoogle Scholar
  2. L.G. Adolfsson, B.A.S. Gustafson, C.D. Murray, The Martian atmosphere as a meteoroid detector. Icarus 199, 144–152 (1996) ADSGoogle Scholar
  3. A.C. Aikin, The lower ionosphere of Mars. Icarus 9, 487–497 (1968). doi:10.1016/0019-1035(68)90042-0 ADSGoogle Scholar
  4. D.E. Anderson, Mariner 6, 7, and 9 ultraviolet spectrometer: Analysis of hydrogen Lyman alpha data. J. Geophys. Res. 79, 1513–1518 (1974) ADSGoogle Scholar
  5. D. Banfield, B.J. Conrath, J. Pearl, M.D. Smith, P.R. Christensen, R.J. Wilson, Forced waves in the Martian atmosphere from MGS/TES nadir data. Icarus 161, 319–345 (2003) ADSGoogle Scholar
  6. C.A. Barth, Photochemistry of the atmosphere of Mars, in The Photochemistry of the Atmospheres, ed. by J.S. Levine (Academic Press, New York, 1985), pp. 337–392 Google Scholar
  7. C.A. Barth, C.W. Hord, Mariner ultraviolet spectrometer topography and polar cap. Science 173, 197–201 (1971). doi:10.1126/science.173.3993.197 ADSGoogle Scholar
  8. C.A. Barth, C.W. Hord, J.B. Pearce, K.K. Kelly, G.P. Anderson, A.I. Stewart, Mariner 6 and 7 ultraviolet spectrometer experiment: Upper atmosphere data. J. Geophys. Res. 76, 2213–2227 (1971) ADSGoogle Scholar
  9. M.J.S. Belton, D.M. Hunten, The abundance and temperature of CO2 in the Martian atmosphere. Astrophys. J. 145, 454–467 (1966) ADSGoogle Scholar
  10. M.J.S. Belton, D.M. Hunten, The spectrographic detection of topographic features on Mars. Science 166, 225–227 (1969) ADSGoogle Scholar
  11. J.-L. Bertaux, F. Leblanc, O. Witasse, E. Quemerais, J. Lilensten, S.A. Stern, B. Sandel, O. Korablev, Discovery of an aurora on Mars. Nature 435, 790–794 (2005) ADSGoogle Scholar
  12. K. Biemann, T. Owen, D.R. Rushneck, A.L. lafleur, D.W. Howarth, The atmosphere of Mars near the surface: Isotope ratios and upper limits on noble gases. Science 194, 76–78 (1976) ADSGoogle Scholar
  13. W.J. Borucki, Z. Levin, R.C. Whitten, R.G. Keesee, L.A. Capone, O.B. Toon, J. Dubach, Predicted electrical conductivity between 0 and 80 km in the Venusian atmosphere. Icarus 51, 302–321 (1982) ADSGoogle Scholar
  14. S.W. Bougher, R.G. Roble, E.C. Ridley, R.E. Dickinson, The Mars thermosphere II. General circulation with coupled dynamical and composition. J. Geophys. Res. 95, 14811–14827 (1990) ADSGoogle Scholar
  15. S.W. Bougher, S. Engel, D.P. Hinson, J.M. Forbes, Mars Global Surveyor radio science electron density profiles: Neutral atmosphere implications. Geophys. Res. Lett. 28(16), 3091–3094 (2001) ADSGoogle Scholar
  16. S.W. Bougher, S. Engel, D.P. Hinson, J.R. Murphy, MGS Radio Science electron density profiles: Interannual variability and implications for the Martian neutral atmosphere. J. Geophys. Res. 109, E03010 (2004) ADSGoogle Scholar
  17. S.W. Bougher, J.R. Murphy, J.M. Bell, M.A. Lopez-Valverde, P.G. Withers, Polar warming in the Mars lower thermosphere: Seasonal variations owing to changing insolation and dust distributions. Geophys. Res. Lett. 33, L02203 (2006) ADSGoogle Scholar
  18. S.W. Bougher, P.-L. Blelly, M. Combi, J.L. Fox, I. Mueller-Wodarg, A. Ridley, R.G. Roble, Neutral upper atmosphere and ionosphere modeling. Space Sci. Rev. 139, 107–141 (2008) ADSGoogle Scholar
  19. S.W. Bougher, T.M. McDunn, K.A. Zoldak, J.M. Forbes, Solar cycle variability of Mars dayside exospheric temperatures: Model evaluation of underlying thermal balance. Geophys. Res. Lett. 36, L05201 (2009). doi:10.1029/2008GL036376 ADSGoogle Scholar
  20. L.H. Brace, H.A. Taylor, T.I. Gombosi, A.J. Kliore, W.C. Knudsen, A.F. Nagy, The ionosphere of Venus: Observations and their interpretations, in Venus, ed. by D.M. Hunten, L. Colin, T.M. Donahue, V.I. Moroz, (1983), pp. 779–840 Google Scholar
  21. D.A. Brain, Mars Global Surveyor measurements of the Martian solar wind interaction. Space Sci. Rev. 126, 77–112 (2006) ADSGoogle Scholar
  22. D. Brain et al., A comparison of global models for the solar wind interaction with Mars. Icarus 206, 149–151 (2010) ADSGoogle Scholar
  23. G.P. Brasseur, S. Solomon, Aeronomy of the Middle Atmosphere, Chemistry and Physics of Stratosphere and Mesosphere (Springer, New York, 2005) Google Scholar
  24. S.H. Brecht, S.A. Ledvina, The solar wind interaction with the Martian ionosphere/atmosphere. Space Sci. Rev. 126, 15–38 (2006) ADSGoogle Scholar
  25. T.K. Breus, A.M. Krymskii, D.H. Crider, N.F. Ness, D. Hinson, K.K. Barashyan, Effect of the solar radiation in the topside atmosphere/ionosphere of Mars: Mars Global Surveyor observations. J. Geophys. Res. 109, A09310 (2004) ADSGoogle Scholar
  26. H.V. Cane, R.E. McGuire, T.T. von Rosenvinge, Two classes of solar energetic particle events associated with impulsive and long-duration soft X-ray flares. Astrophys. J. 301, 448–459 (1986) ADSGoogle Scholar
  27. T. Cavalie et al., Vertical temperature profile and mesospheric winds retrieval on Mars from CO; millimeter observations. Comparison with general circulation model predictions. Astron. Astrophys. 489, 795–809 (2008) ADSGoogle Scholar
  28. R.H. Chen, T.E. Cravens, A.F. Nagy, The Martian ionosphere in light of the Viking observations. J. Geophys. Res. 83, 3871–3876 (1978) ADSGoogle Scholar
  29. A.A. Christou, Annual meteor showers at Venus and Mars: Lession from the Earth. Mon. Not. R. Astron. Soc. 402, 2759–2770 (2010) ADSGoogle Scholar
  30. A.A. Christou, K. Beurle, Meteoroid streams at Mars: Possibilities and implications. Planet. Space Sci. 47, 1475–1485 (1999) ADSGoogle Scholar
  31. F. Cipriani, F. Leblanc, J.J. Berthelier, Martian corona: Nonthermal sources of hot heavy species. J. Geophys. Res. 112, E07001 (2007). doi:10.1029/2006JE002818 ADSGoogle Scholar
  32. R.T. Clancy, M.J. Wolff, P.B. James, Minimal aerosols loading and global increases in atmospheric ozone during 1996–1997 Martian northern spring seasons. Icarus 138, 49–63 (1999) ADSGoogle Scholar
  33. T.E. Cravens, A.F. Nagy, Aeronomy of inner planets. Rev. Geophys. 21, 263–273 (1983). doi:10.1029/RG021i002p00263 ADSGoogle Scholar
  34. J.A. Crisp, M. Adler, J.R. Matijevic, S.W. Squyres, R.E. Arvidson, D.M. Kass, Mars exploration rover mission. J. Geophys. Res. 108(E12), 8061–8071 (2003). doi:10.1029/2002JE002038 Google Scholar
  35. A. Dalgarno, M.B. McElroy, Mars: Is nitrogen present? Science 170, 167–168 (1970) ADSGoogle Scholar
  36. I. Dandouras, H. Reme, J.B. Cao, P. Escoubet, P.C. Brandt, Abstract on magnetosphere response to the 2005 and 2006 extreme solar events as observed by the Cluster and Double Star Spacecraft: Solar Extreme Events. Symposium held at Athens in September 2007 Google Scholar
  37. A. Domokos, J.F. Bell, P. Brown, M.T. Lemmon, R. Suggs, J. Vaubaillon, W. Cook, Measurement of the meteoroid flux at Mars. Icarus 191, 141–150 (2007) ADSGoogle Scholar
  38. F. Duru, D.A. Gurnett, D.D. Morgan, R. Modolo, A.F. Nagy, D. Najib, Electron densities in the upper ionosphere of Mars from the excitation of electron plasma oscillations. J. Geophys. Res. 113, A07302 (2008) ADSGoogle Scholar
  39. F. Duru et al., Steep, transient densitygradients in the Martian ionosphere similar to the ionopause at Venus. J. Geophys. Res. 114, A12310 (2009). doi:10.1029/2009JA014711 ADSGoogle Scholar
  40. X. Fang, M.W. Liemohn, A.F. Nagy, J. Luhmann, Y. Ma, On the effect of the Martian crustal magnetic field on atmospheric erosion. Icarus 206, 130–138 (2010) ADSGoogle Scholar
  41. K. Fast et al., Ozone abundance on Mars from infrared heterodyne spectra II. Validating photochemical models. Icarus 183, 396–402 (2006) ADSGoogle Scholar
  42. M.O. Fillingim, L.M. Peticolas, R.J. Lillis, D.A. Brain, J.S. Halekas, D. Lummerzheim, S.W. Bougher, Localized ionization patches in the nighttime ionosphere of Mars and their electrodynamic consequences. Icarus 206, 112–119 (2010) ADSGoogle Scholar
  43. G. Fjeldbo, W.C. Fjeldbo, V. Eshleman, Models for the atmosphere of Mars based on the Mariner 4 occultation experiment. J. Geophys. Res. 71, 2307–2316 (1966) ADSGoogle Scholar
  44. G. Fjeldbo, A. Kliore, B. Seidel, The Martian 1969 occultation measurements of the upper atmosphere of Mars. Radio Sci. 5, 381–386 (1970) ADSGoogle Scholar
  45. G. Fjeldbo, D. Sweetnam, J. Brenkle, E. Christensen, D. Farless, J. Mehta, B. Seidel, W. Michael Jr., A. Wallio, M. Grossi, Viking radio occultation measurements of Martian atmosphere and topography: Primary mission covering age. J. Geophys. Res. 82, 4317–4324 (1977) ADSGoogle Scholar
  46. J.L. Fox, Airglow and Aurora in the atmosphere of Venus and Mars, in Venus and Mars: Atmosphere, Ionosphere, and Solar Wind Interactions, ed. by J.G. Luhmann, M. Tatrallyay, R.O. Pepin. Geophys. Monogr. Ser., vol. 66 (AGU, Washington, 1992), pp. 191–222 Google Scholar
  47. J.L. Fox, The production and escape of nitrogen atoms on Mars. J. Geophys. Res. 98, 3297–3310 (1993). doi:10.1029/92JE02289 ADSGoogle Scholar
  48. J.L. Fox, Morphology of the dayside ionosphere of Mars: Implications for ion outflows. J. Geophys. Res. 114, E12005 (2009). doi:10.1029/2009JE003432 ADSGoogle Scholar
  49. J.L. Fox, A. Dalgarno, Ionization, luminosity, and heating of the upper atmosphere of Mars. J. Geophys. Res. 84, 7315–7331 (1979) ADSGoogle Scholar
  50. J.L. Fox, A. Hac, Spectrum of hot O at the exobase of the terrestrial planets. J. Geophys. Res. 102, 24005–24011 (1997). doi:10.1029/97JA02089 ADSGoogle Scholar
  51. J.L. Fox, A.J. Weber, MGS electron density profiles: Analysis and modelling of peak altitudes. Icarus 221, 1002–1019 (2012) ADSGoogle Scholar
  52. J.L. Fox, K.E. Yeager, Morphology of the near termination Martian ionosphere: A comparison of models and data. J. Geophys. Res. 111, A10309 (2006) ADSGoogle Scholar
  53. J.L. Fox, K.E. Yeager, MGS electron density profiles: Analysis of the peak magnitudes. Icarus 200, 468–479 (2009) ADSGoogle Scholar
  54. J.L. Fox, J.F. Brannon, H.S. Porter, Upper limits to the nightside ionosphere of Mars. Geophys. Res. Lett. 20, 1339–1342 (1993) ADSGoogle Scholar
  55. J.L. Fox, P. Zhon, S.W. Bougher, The Martian thermosphere/ionosphere at high and low solar activities. Adv. Space Res. 17(11), 203 (1996) ADSGoogle Scholar
  56. D. Grassi et al., The Martian atmosphere above great volcanoes: Early Planetary Fourier Spectrometer observations. Planet. Space Sci. 53, 1053–1064 (2005) ADSGoogle Scholar
  57. D.A. Gurnett et al., Radar soundings of the ionosphere of Mars. Science 310, 1929–1933 (2005) ADSGoogle Scholar
  58. D.A. Gurnett et al., An overview of radar soundings of the Martian ionosphere from the Mars Express spacecraft. Adv. Space Res. 41, 1335–1346 (2008). doi:10.1016/j.asr.2007.01.062 ADSGoogle Scholar
  59. S.A. Haider, Chemistry on the nightside ionosphere of Mars. J. Geophys. Res. 102, 407–416 (1997). doi:10.1029/96JA02353 ADSGoogle Scholar
  60. S.A. Haider, Role of X-ray flares and CME in the E region ionosphere of Mars: MGS observations. Planet. Space Sci. 63/64, 56–61 (2012) ADSGoogle Scholar
  61. S.A. Haider, J. Kim, A.F. Nagy, C.N. Keller, M.I. Verigin, K.I. Gringauz, N.M. Shutte, K. Szego, P. Kiraly, Calculated ionization rates, ion densities, and airglow emission rates due to precipitating electrons in the nightside ionosphere of Mars. J. Geophys. Res. 97(A7), 10637–10641 (1992). doi:10.1029/92JA00317 ADSGoogle Scholar
  62. S.A. Haider, S.P. Seth, E. Kallio, K.I. Oyama, Solar EUV and electron-proton-hydrogen atom produced ionosphere on Mars: Comparative studies of particle fluxes and ion production rates due to different processes. Icarus 159, 18–30 (2002). doi:10.1006/icar.2002.6919 ADSGoogle Scholar
  63. S.A. Haider, S.P. Seth, V.R. Choksi, K.I. Oyama, Model of photoelectron impact ionization within the high latitude ionosphere at Mars: Comparison of calculated and measured electron density. Icarus 185, 102–112 (2006). doi:10.1016/j.icarus.2006.07.010 ADSGoogle Scholar
  64. S.A. Haider, V. Singh, V.R. Choksi, W.C. Maguire, M.I. Verigin, Calculated densities of H3O+(H2O)n, \(\mathrm{NO}_{2}^{-}(\mathrm{H}_{2}\mathrm{O})_{\mathrm{n}}\), \(\mathrm{CO}_{3}^{-}(\mathrm{H}_{2}\mathrm{O})_{\mathrm{n}}\) and electron in the nighttime ionosphere of mars: impact of solar wind electron and galactic cosmic rays. J. Geophys. Res. 112, A12309 (2007). doi:10.1029/2007JA012530 ADSGoogle Scholar
  65. S.A. Haider, V. Sheel, V. Singh, W.C. Maguire, G.J. Molina-Cuberos, Model calculation of production rates, ion and electron densities in the evening troposphere of Mars at altitudes 67N and 62S: Seasonal variability. J. Geophys. Res. 113, A08320 (2008). doi:10.1029/2007JA012980 ADSGoogle Scholar
  66. S.A. Haider, M.A. Abdu, I.S. Batista, J.H. Sobral, E. Kallio, E. Kallio, W.C. Maguire, M.I. Verigin, On the responses to solar X-ray flare and coronal mass ejection in the ionosphere of Mars and Earth. Geophys. Res. Lett. 36, L13104 (2009a). doi:10.1029/2009GL038694 ADSGoogle Scholar
  67. S.A. Haider, M.A. Abdu, I.S. Batista, J.H. Sobral, X. Luan, E. Kallio, W.C. Maguire, M.I. Verigin, V. Singh, D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: Comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. 114, A03311 (2009b). doi:10.1029/2008JA13709 ADSGoogle Scholar
  68. S.A. Haider, M.A. Abdu, I.S. Batista, J.H. Sobral, V. Sheel, G.J. Molina-Cuberos, W.C. Magurie, M.I. Verigin, Zonal wave structures in the nighttime tropospheric density and temperature and in the D-region ionosphere over Mars: Modeling and observation. J. Geophys. Res. 114, A12351 (2009c) Google Scholar
  69. S.A. Haider, S.P. Seth, D.A. Brain, D.L. Mitchell, T. Majeed, S.W. Bougher, Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis Major: MGS observations. J. Geophys. Res. 115, A08310 (2010a). doi:10.1029/2009JA014968 ADSGoogle Scholar
  70. S.A. Haider, V. Sheel, M.D. Smith, W.C. Maguire, G.J. Molina-Cuberos, Effect of dust storms on the D region of the Martian Ionosphere: Atmospheric electricity. J. Geophys. Res. 115, A12336 (2010b). doi:10.1029/2010JA016125 ADSGoogle Scholar
  71. S.A. Haider, K.K. Mahajan, E. Kallio, Mars ionosphere: A review of experimental results and modeling studies. Rev. Geophys. 49, RG4001 (2011) ADSGoogle Scholar
  72. S.A. Haider, S.M.P. McKenna-Lawlor, C.D. Fry, R. Jain, K.N. Joshipura, Effects of solar X-ray flares in the E region ionosphere of Mars: First model results. J. Geophys. Res. 117, A05326 (2012). doi:10.1029/2011JA017436 ADSGoogle Scholar
  73. S.A. Haider, B.M. Pandya, G.J. Molina-Cuberos, Nighttime ionosphere caused by meteoroid ablation and solar wind electron-proton-hydrogen impact: MEX observation and modelling. J. Geophys. Res. 115, 1–9 (2013). doi:10.1002/jgra.50590 Google Scholar
  74. W.B. Hanson, G.P. Mantas, Viking electron temperature measurements: Evidence for a magnetic field in the Martian atmosphere. J. Geophys. Res. 93, 7538–7544 (1988) ADSGoogle Scholar
  75. W.B. Hanson, S. Sanatani, R. Zuccaro, The Martian ionosphere as observed by the Viking retarding potential analyzers. J. Geophys. Res. 82, 4351–4363 (1977) ADSGoogle Scholar
  76. J.K. Hargreaves, The Solar-Terrestrial Environment: An Introduction to Geospace—The Science of Terrestrial Upper Atmosphere, Ionosphere and Magnetosphere (Cambridge Univ. Press, New York, 1992) Google Scholar
  77. M.G. Heavens et al., Vertical distribution of dust in the Martian atmosphere during northern spring and summer: High altitude tropical dust maximum at northernsummer solstice. J. Geophys. Res. 116, E01007 (2011). doi:10.1029/2010JE003692 ADSGoogle Scholar
  78. S.L. Hess et al., Mars climatology from Viking 1 after 20 sols. Science 194, 78–81 (1976a) ADSGoogle Scholar
  79. S.L. Hess et al., Early meteorological results from Viking 2 Lander. Science 194, 1352–1353 (1976b) ADSGoogle Scholar
  80. D.P. Hinson, Mars global surveyor radio occultation profiles of the ionosphere-reorganized, MGS-M-RSS-5-EDS-V1.0, vol. USA_NASA_JPL_MORS_1102, NASA Planetary Data System (NASA Goddard Space Flight Center, Greenbelt, 2007) Google Scholar
  81. D.P. Hinson, R.J. Wilson, Temperature inversion, thermal tides, and water ice clouds in the Martian tropics. J. Geophys. Res. 109, E01002 (2004). doi:10.1029/2003JE00129 ADSGoogle Scholar
  82. D.P. Hinson, R.A. Simpson, J.D. Twicken, G.L. Tyler, F.M. Flassar, Initial results from radio occultation measurements with Mars Global Surveyor. J. Geophys. Res. 104, 26997–27012 (1999) ADSGoogle Scholar
  83. D.P. Hinson, M.D. Smoth, B.J. Conrath, Comparison of atmospheric temperatures obtained through infrared sounding and radio occultation by Mars Global Surveyor. J. Geophys. Res. 109, E12002 (2004). doi:10.1029/2004JE002344 ADSGoogle Scholar
  84. J.L. Hollingsworth, J.R. Barnes, Forced, stationary planetary waves in Mars’winter atmosphere. J. Atmos. Sci. 53, 428–448 (1996) ADSGoogle Scholar
  85. F. Hourdin, P. LeVan, F. Forget, O. Talagrand, Meteorological variability and annual surface pressure cycle on Mars. J. Atmos. Sci. 50, 3625–3640 (1993) ADSGoogle Scholar
  86. D.M. Hunten, Escape of atmospheres, ancient and modern. Icarus 85, 1–20 (1990) ADSGoogle Scholar
  87. T. Imamura, T. Ogawa, Radiative damping og gravity waves in the terrestrial planetary atmospheres. Geophys. Res. Lett. 22, 267–270 (1995) ADSGoogle Scholar
  88. W.H. Ip, Meteoroid ablation processes in Titan’s atmosphere. Nature 345, 511–512 (1990) ADSGoogle Scholar
  89. W.H. Ip, ENA diagnostic of auroral activity at Mars. Planet. Space Sci. 63/64, 83–86 (2012) ADSGoogle Scholar
  90. V.G. Istomin, K.V. Grechnev, Argon in the Martian atmosphere: Evidence from the Mars 6 descent module. Icarus 28, 155–158 (1976) ADSGoogle Scholar
  91. E. Kallio, P. Janhunen, Atmospheric effects of proton precipitation in the Martian atmosphere and its connection to the Mars-solar wind interaction. J. Geophys. Res. 106, 5617–5634 (2001) ADSGoogle Scholar
  92. E. Kallio, K. Liu, R. Javinen, V.Pohjola.P. Janhunen, Oxygen ion excape at Mars in ahybrid model: High energy and low energy ions. Icarus 206, 152–163 (2010). doi:10.1016/j.icarus.2009.05.015 ADSGoogle Scholar
  93. L.D. Kaplan, J. Connes, P. Connes, Carbon monoxide in the Martian atmosphere. Astrophys. J. 157, L187–L192 (1969) ADSGoogle Scholar
  94. J. Kar, Recent advances in planetary ionospheres. Space Sci. Rev. 77, 193–266 (1996). doi:10.1007/BF00226224 ADSGoogle Scholar
  95. G.M. Keating et al., The structure of the upper atmosphere of Mars: In situ accelerometer measurements from Mars Global Surveyor. Science 279, 1672–1675 (1998). doi:10.1126/science.279.5357.1672 ADSGoogle Scholar
  96. M.C. Kelley, The Earth’s Ionosphere (Academic Press, San Diego, 1989) Google Scholar
  97. J. Kim, A.F. Nagy, J.L. Fox, T.E. Cravens, Solar cycle variability of hot oxygen atoms at Mars. J. Geophys. Res. 103, 29339–29342 (1998) ADSGoogle Scholar
  98. A.J. Kliore, J.G. Luhmann, Solar cycle effects on the structure of the electron density profiles in the dayside ionosphere of Venus. J. Geophys. Res. 96, 21281–21289 (1991) ADSGoogle Scholar
  99. A.J. Kliore, D.L. Cane, G.S. Levy, V.R. Eshleman, G. Fjeldbo, F.D. Drake, Occultation experiment: Results of the first direct measurement of Mars’ atmosphere and ionosphere. Science 149, 1243–1248 (1965) ADSGoogle Scholar
  100. A.J. Kliore, D.L. Cain, G. Fjeldbo, B.L. Seidel, M.J. Sykes, S.I. Rasool, The atmosphere of Mars from Mariner 9 radio occultation measurements. Icarus 17, 484–516 (1972) ADSGoogle Scholar
  101. A.J. Kliore, G. Fjeldbo, B.L. Seidel, M.J. Sykes, P.M. Woiceshyn, S band radio occultation measurements of the atmosphere and topography of Mars with Mariner 9: Extended Mission coverage of polar and intermediate latitudes. J. Geophys. Res. 78, 4331–4351 (1973) ADSGoogle Scholar
  102. M.A. Kolosov, O.I. Yakovlev, Yu.M. Kruglov, B.P. Trusov, A.I. Efimov, V.V. Kerzhanovich, Radio sounding of the Martian atmosphere by spacecraft. Radio Eng. Electron. Phys. 17, 1993–1999 (1972) Google Scholar
  103. M.A. Kolosov et al., Results of two-frequency radio occultation of ‘Mars-2’ by Ionosphere of Mars. Radio Eng. Electron. Phys. 18, 1471–1474 (1973) Google Scholar
  104. M.A. Kolosov et al., Results of investigating the Martian atmosphere by radio occultation using Mars 2, Mars 4 and Mars 6 spacecraft. Kosm. Issled 13, 54–59 (1975) ADSGoogle Scholar
  105. T.Y. Kong, M.B. McElroy, Photochemistry of the Martian atmosphere. Icarus 32, 168–176 (1977) ADSGoogle Scholar
  106. A.J. Kopf, D.A. Gurnett, D.D. Morgan, D.L. Kirchner, Transient layers in the topside ionosphere of Mars. Geophys. Res. Lett. L17102 (2008). doi:10.1029/2008GL034948
  107. E. Kopp, U. Hermann, Ion composition in the lower ionosphere. Ann. Geophys. 2, 83–94 (1984) ADSGoogle Scholar
  108. V.A. Krasnopolsky, Mars’ upper atmosphere and ionosphere at low, medium, and high solar activities: implications for evolution of water. J. Geophys. Res. 107, 5128 (2002) Google Scholar
  109. V.A. Krasnopolsky, Spectroscopic mapping of Mars CO mixing ratio: Detection of north south asymmetry. J. Geophys. Res. 108(E2), 5010 (2003a) Google Scholar
  110. V.A. Krasnopolsky, Mapping of Mars O2 1.27 mm dayglow at four seasonal points. Icarus 165, 315–325 (2003b) ADSGoogle Scholar
  111. V.A. Krasnopolsky, V.A. Parshev, Ozone photochemistry of the Martian lower atmosphere. Planet. Space Sci. 27, 113–120 (1979) ADSGoogle Scholar
  112. V.A. Krasnopolsky, S. Bowyer, S. Chakrabarti, G.R. Gladstone, J.S. McDonald, First measurement of helium on Mars: Implications for the problem of radiogenic gases on the terrestrial planets. Icarus 109, 337–351 (1994). doi:10.1006/icar.1994.1098 ADSGoogle Scholar
  113. A.M. Krymskii, T.K. Breus, N.F. Ness, D.P. Hinson, D.I. Bojkov, Effect of crustal magnetic fields on the near terminator ionospheres at Mars: Comparison of in situ magnetic field measurements with the data of radio science experiments on board Mars Global Surveyor. J. Geophys. Res. 108, 1431 (2003). doi:10.1029/2002JA009662 Google Scholar
  114. G.P. Kuiper (ed.), The Atmosphere of Earth and Planets (1952). Chapter XII Google Scholar
  115. S. Kumar, D.M. Hunten, Venus: An ionospheric model with an exospheric temperature of 350 K. J. Geophys. Res. 79, 2529–2532 (1974) ADSGoogle Scholar
  116. A. Kumar, N.K. Lodhi, K.K. Mahajan, Near terminator ionosphere during sunspot cycle 23 from Mars Global Surveyor radio science measurements. Indian J. Radio Space Phys. 36, 457–465 (2007) Google Scholar
  117. H. Lammer, S.J. Bauer, Nonthermal atmospheric escape from Mars and Titan. J. Geophys. Res. 96, 1819–1825 (1991). doi:10.1029/90ja01676 ADSGoogle Scholar
  118. H. Lammer, W. Stumptner, S.J. Bauer, Upper limit for the Martian exospheric number density during the Planet B/Nozomi mission. Planet. Space Sci. 48, 1473–1478 (2000) ADSGoogle Scholar
  119. F. Leblanc, J.G. Luhmann, R.E. Johnson, E. Chassefiere, Some expected impacts of a solar energetic particle event at Mars. J. Geophys. Res. 107, 1058 (2002) Google Scholar
  120. F. Leblanc et al., Observations of aurorae by SPICAM ultraviolet spectrograph on board Mars Express: Simultaneous ASPERA-3 and MARSIS measurements. J. Geophys. Res. 113, A08311 (2008) ADSGoogle Scholar
  121. S.A. Ledvina, Y.-J. Ma, E. Kallio, Modeling and simulating flowing plasmas and related phenomena. Space Sci. Rev. 139, 143–189 (2008). doi:10.1007/s11214-008-9384-6 ADSGoogle Scholar
  122. F. Lefevre, S. Lebonnois, F. Montmessin, F. Forget, Three dimensional modeling of ozone on Mars. J. Geophys. Res. 109, E07004 (2004). doi:10.1029/2004JE002268 ADSGoogle Scholar
  123. F. Lefevre et al., Heterogeneous chemistry in the atmosphere of Mars. Nature 454, 971–975 (2008) ADSGoogle Scholar
  124. L. Lei, Y. Zhang, Model investigation of the influence of the crustal magnetic field on the oxygen ion distribution in the near Martian tail. J. Geophys. Res. 114, A06215 (2009). doi:10.1029/2008JA013850 ADSGoogle Scholar
  125. R.J. Lillis, D.L. Mitchell, R.P. Lin, M.H. Acuna, Electron reflectometry in the Martian atmosphere. Icarus 194, 544–561 (2008). doi:10.1016/j.icarus.2007.09.030 ADSGoogle Scholar
  126. R.J. Lillis, M.O. Fillingim, L.M. Peticolas, D.A. Brain, R.P. Lin, S.W. Bougher, The nightside ionosphere of Mars: Modeling the effects of crustal magnetic fields and electron pitch angle distributions on electron impact ionization. J. Geophys. Res. 114, E11009 (2009) ADSGoogle Scholar
  127. R.J. Lillis, D.A. Brain, S.L. England, P. Withers, M.O. Fillingim, A. Safaeinili, Total electron content in the Mars ionosphere: Temporal studies and dependence on solar EUV flux. J. Geophys. Res. 115, A11314 (2010) ADSGoogle Scholar
  128. R.J. Lillis, M.O. Fillingim, D.A. Brain, Three-dimensional structure of the Martian nightside ionosphere: Predicted rates of impact ionization from Mars Global Surveyor magnetometer and electron reflectometer measurements of precipitating electrons. J. Geophys. Res. 116, A12317 (2011) ADSGoogle Scholar
  129. G.F. Lindal, H.B. Hotz, D.N. Sweetnam, Z. Shippony, J.P. Brenkle, G.V. Hartsell, R.T. Spear, W.H. Michael Jr., Viking radio occultation measurements of the atmosphere and topography of Mars: Data acquired during 1 Martian year of tracking. J. Geophys. Res. 84, 8443–8456 (1979) ADSGoogle Scholar
  130. J.L. Lovell, M.L. Dulding, J.E. Humble, An extended analysis of the September 1989 Cosmic ray ground level enhancement. J. Geophys. Res. 103, 23733–23742 (1998). doi:10.1029/98JA02100 ADSGoogle Scholar
  131. R. Lundin et al., Plasma acceleration above Martian magnetic anomalies. Science 311, 980 (2006). doi:10.1126/science.1122071 ADSGoogle Scholar
  132. R. Lundin, S. Barabash, E. Dubinin, D. Winingham, M. Yamauchi, Low latitudeacceleration of ionospheric ions at Mars. Geophys. Res. Lett. 38, L08108 (2011). doi:10.1029/2011GL047064 ADSGoogle Scholar
  133. Y. Ma, A.F. Nagy, I.V. Sokolov, K.C. Hanse, Three dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. 109, A07211 (2004) ADSGoogle Scholar
  134. Y. Ma et al., Plasma flow and related phenomena in planetary aeronomy. Space Sci. Rev. 139, 311–353 (2008). doi:10.1007/s11214-008-9389-1 ADSGoogle Scholar
  135. J.A. Magalhaes, J.T. Schofield, A. Seiff, Results of the Mars Pathfinder atmospheric structure investigation. J. Geophys. Res. 104(E4), 8943–8955 (1999) ADSGoogle Scholar
  136. K.K. Mahajan, J. Kar, Planetary ionosphere. Space Sci. Rev. 47, 303–397 (1988). doi:10.1007/BF00243558 ADSGoogle Scholar
  137. K.K. Mahajan, S. Singh, A. Kumar, S. Raghuvanshi, S.A. Haider, Mars Global Surveyor radio science electron density profiles: Some anomalous features in the Martian ionosphere. J. Geophys. Res. 112, E10006 (2007). doi:10.1029/2006JE002876 ADSGoogle Scholar
  138. K.K. Mahajan, N.K. Lodhi, S. Singh, Ionospheric effects of solar flares at Mars. Geophys. Res. Lett. 36, L15207 (2009). doi:10.1029/2009GL039454 ADSGoogle Scholar
  139. A.P. Mayo et al., Lander locations, Mars physical ephemeris, and solar system parameters: Determination from Viking Lander tracking data. J. Geophys. Res. 82, 4297–4303 (1977) ADSGoogle Scholar
  140. M.B. McElroy, The upper atmosphere of Mars. Astrophys. J. 150, 1125–1138 (1967) ADSGoogle Scholar
  141. M.B. McElroy, T.M. Donahue, Stability of the Martian atmosphere. Science 177, 986–988 (1972) ADSGoogle Scholar
  142. M.B. McElroy, Y.L. Yung, A.O. Neir, Isotopic composition of nitrogen: Implications for the past history of Mars atmosphere. Science 194, 70–72 (1976) ADSGoogle Scholar
  143. S. McKenna-Lawlor, P. Goncalves, A. Keating, G. Reitz, D. Matthia, Overview of energetic particle hazards during prospective manned mission to Mars. Planet. Space Sci. 63/64, 12–132 (2012) Google Scholar
  144. M. Mendillo, S. Smith, J. Wroten, H. Rishbeth, D. Hinson, Simultaneous ionospheric variability on Earth and Mars. J. Geophys. Res. 108, 1432–1443 (2003) Google Scholar
  145. M. Mendillo, X. Pi, S. Smith, C. Martinis, J. Wilson, D. Hinson, Ionospheric effects upon a satellite navigation system at Mars. Radio Sci. 39, RS2028 (2004) ADSGoogle Scholar
  146. M. Mendillo, P. Withers, D. Hinson, H. Rishbeth, B. Reinisch, Effects of solar flares on the ionosphere of Mars. Science 311, 1135–1138 (2006) ADSGoogle Scholar
  147. M. Mendillo, A. Lollo, P. Withers, M. Matta, M. Pätzold, S. Tellmann, Modeling Mars’ ionosphere with constraints from same-day observations by Mars Global Surveyor and Mars Express. J. Geophys. Res. 116, A11303 (2011) ADSGoogle Scholar
  148. M. Michael, M. Barani, S.N. Tripathi, Numerical predictions of aerosol charging and electrical conductivity of the lower atmosphere of Mars. Geophys. Res. Lett. 34, L04201 (2007). doi:10.1029/2006GL028434 ADSGoogle Scholar
  149. D.L. Mitchell, R.P. Lin, H. Rème, D.H. Crider, P.A. Cloutier, J.E.P. Connerney, M.H. Acuña, N.F. Ness, Oxygen auger electrons observed in Mars ionosphere. Geophys. Res. Lett. 27, 1871–1874 (2000) ADSGoogle Scholar
  150. D.L. Mitchell, R.J. Lillis, R.P. Lin, J.E.P. Connerney, M.H. Acuña, A global map of Mars’ crustal magnetic field based on electron reflectometer. J. Geophys. Res. 112, E01002 (2007) ADSGoogle Scholar
  151. R. Modolo, G.M. Chanteur, E. Dubinin, A.P. Matthews, Simulated solar wind plasma interaction with the Martian exosphere: Influence of the solar EUV flux on the bow shock and the magnetic pile-up boundary. Ann. Geophys. 24, 3403–3410 (2006) ADSGoogle Scholar
  152. G.J. Molina-Cuberos, H. Lichtenegger, K. Schwingenschuh, J.J. Lopez-Moreno, R. Rodrigo, Ion-neutral chemistry model of the lower ionosphere of Mars. J. Geophys. Res. 105(E5), 31–37 (2002) Google Scholar
  153. G.J. Molina-Cuberos, O. Witasse, J.-P. Lebreton, R. Rodrigo, J.J. Loṕez-Moreno, Meteoric ions in the atmosphere of Mars. Planet. Space Sci. 51, 239–249 (2003) ADSGoogle Scholar
  154. G.J. Molina-Cuberos, J.J. López-Moreno, F. Arnold, Meteoric Layers in planetary atmospheres. Space Sci. Rev. 137, 175–191 (2008) ADSGoogle Scholar
  155. L. Montabone, S.R. Lewis, P.L. Read, D.P. Hinson, Validation of Martian meteorological data assimilation for MGS/TES using radio occultation measurements. Icarus 185, 113–132 (2006) ADSGoogle Scholar
  156. D.D. Morgan, D.A. Gurnett, D.L. Kirchner, R.L. Huff, D.A. Brain, W.V. Boynton, M.H. Acuña, J.J. Plaut, G. Picardi, Solar control of radar wave absorption by the Martian ionosphere. Geophys. Res. Lett. 33, L13202 (2006) ADSGoogle Scholar
  157. D.D. Morgan, D.A. Gurnett, D.L. Kirchner, J.L. Fox, E. Nielsen, J.J. Plaut, Variation of the Martian ionospheric electron density from Mars Express radar soundings. J. Geophys. Res. 113, A09303 (2008) ADSGoogle Scholar
  158. V.I. Moroz, The atmosphere of Mars. Space Sci. Rev. 19, 763–843 (1976) ADSGoogle Scholar
  159. P.M. Mul, J.W. McGowan, Temperature dependence of dissociative recombination for atmospheric ions NO, O2, N2. J. Phys. B 12, 1591–1602 (1979) ADSGoogle Scholar
  160. A.F. Nagy, T.E. Cravens, S.G. Smith, H.A. Taylor, H.C. Brinton, Model calculations of the dayside ionosphere of Venus-Ionic composition. J. Geophys. Res. 85, 7795–7801 (1980) ADSGoogle Scholar
  161. A.F. Nagy, M.W. Liemohn, J.L. Fox, J. Kim, Hot carbon densities in the exosphere of Mars. J. Geophys. Res. 106, 21565–21572 (2001) ADSGoogle Scholar
  162. A.F. Nagy et al., The plasma environment of Mars. Space Sci. Rev. 111, 33–114 (2004) ADSGoogle Scholar
  163. H. Nair, M. Allen, A.D. Anbar, Y.L. Yung, R.T. Claney, A photochemical model of the Martian atmosphere. Icarus 111, 124–150 (1994) ADSGoogle Scholar
  164. N.F. Ness, M.H. Acuna, J.E.P. Connerney, A.J. Kliore, T.K. Breus, A.M. Krymskii, P. Cloutier, S.J. Bauer, Effects of magnetic anomalies discovered at Mars on the structure of the Martian ionosphere and solar wind interaction as follows from radio occultation experiments. J. Geophys. Res. 105, 15991–16004 (2000) ADSGoogle Scholar
  165. E. Nielsen, H. Zou, D.A. Gurnett, D.L. Kirchner, D.D. Morgan, R. Huff, R. Orosei, A. Safaeinili, J.J. Plaut, G. Picardi, Observations of vertical reflections from the topside martian ionosphere. Space Sci. Rev. 126, 373–388 (2006) ADSGoogle Scholar
  166. A.O. Nier, M.B. McElroy, Composition and structure of Mars’ upper atmosphere: Results from the neutral mass spectrometers on Viking 1 and 2. J. Geophys. Res. 82, 4341–4349 (1977) ADSGoogle Scholar
  167. K. O’Brien, K.W. Friedberg, H.H. Sauer, D.F. Smart, Atmospheric cosmic rays and solar energetic particles at aircraft altitudes. Environ. Int. 22, S9–S44 (1996) Google Scholar
  168. T. Owen, K. Biemann, D.R. Rushneck, J.E. Biller, D.W. Howarth, A.L. Lafleur, The atmosphere of Mars: Detection of Krypton and Xenon. Science 194, 1293–1295 (1976) ADSGoogle Scholar
  169. T. Owen, K. Biemann, D.R. Rushneck, J.E. Biller, D.W. Howarth, A.L. Lafleur, The composition of the atmosphere at the surface of Mars. J. Geophys. Res. 82, 4635–4644 (1977) ADSGoogle Scholar
  170. B.M. Pandya, S.A. Haider, Meteor impact perturbation in the lower ionosphere of Mars: MGS observations. Planet. Space Sci. 63/64, 105–109 (2012) ADSGoogle Scholar
  171. T.D. Parkinson, D.M. Hunten, Spectroscopy and aeronomy of O2 on Mars. J. Atmos. Sci. 29, 1380–1390 (1972) ADSGoogle Scholar
  172. M. Pätzold, S. Tellmann, B. Ha¨usler, D. Hinson, R. Schaa, G.L. Tyler, A sporadic third layer in the ionosphere of Mars. Science 310, 837–839 (2005) ADSGoogle Scholar
  173. J. Perrier et al., Global distribution of total ozone on Mars from SPICAM/MEX UV measurements. J. Geophys. Res. 111, E09S06 (2006). doi:10.1029/2006JE002681 ADSGoogle Scholar
  174. W.D. Pesnell, J.M. Grebowsky, Meteoric Magnesium in the Martian atmosphere. J. Geophys. Res. 105, 1695–1703 (2000) ADSGoogle Scholar
  175. G. Picardi et al., MARSIS: Mars advanced radar for subsurface and ionospheric sounding, in Mars Express: The Scientific Payload, ed. by A. Wilson, A. Chicarro. Eur. Space Agency Spec. Publ., ESA-SP, vol. 1240 (2004), pp. 51–59 Google Scholar
  176. S.I. Rasool, R.W. Stewart, Results and interpretation of the S-band occultation experiment on Mars and Venus. J. Atmos. Sci. 28, 869–878 (1971) ADSGoogle Scholar
  177. S.I. Rasool, J.S. Hogan, R.W. Stewart, L.H. Russell, Temperature distributions in the lower atmosphere of Mars from Mariner 6 and 7 radio occultation data. J. Atmos. Sci. 27, 841–843 (1970) ADSGoogle Scholar
  178. H. Rishbeth, O.K. Garriott, Introduction to Ionospheric Physics (Elsevier, New York, 1969) Google Scholar
  179. H. Rishbeth, M. Mendillo, Patterns of F2 layer variability. J. Atmos. Sol.-Terr. Phys. 63, 1661–1680 (2001) ADSGoogle Scholar
  180. H. Rishbeth, M. Mendillo, Ionospheric layers of Mars and Earth. Planet. Space Sci. 52, 849–852 (2004) ADSGoogle Scholar
  181. R. Rodrigo, E. Gracia-Alvarez, M.J. Lopez-Gonzalez, J.J. Lopez-Moreno, A non-steady one dimensional theoretical model of Mars’ neutral atmospheric composition between 30 and 200 km. J. Geophys. Res. 95, 14795–14810 (1990) ADSGoogle Scholar
  182. R.P. Rohrbaugh, J.S. Nisbet, E. Bleuler, J.R. Herman, The effects of energetically produced \(\mathrm{O}^{+}_{2}\) on the ion temperature of the Martian thermosphere. J. Geophys. Res. 84, 3327–3336 (1979) ADSGoogle Scholar
  183. A. Safaeinili, W. Kofman, J. Mouginot, Y. Gim, A. Herique, A.B. Ivanov, J.J. Plaut, G. Picardi, Estimation of the total electron content of the Martian ionosphere using radar sounder surface echoes. Geophys. Res. Lett. 34, L23204 (2007) ADSGoogle Scholar
  184. N.A. Savich, V.A. Samovol, The night time ionosphere of Mars from Mars 4 and Mars 5 dual frequency radio occultation measurements. Space Res. XVI, 1009–1010 (1976) ADSGoogle Scholar
  185. C.J. Schnjver, G.L. Siscoe, Heliophysics (Cambridge Univ. Press, Cambridge, 2010) Google Scholar
  186. R.W. Schunk, A.F. Nagy, Ionosphere of the terrestrial planets. Rev. Geophys. 18, 813–852 (1980). doi:10.1029/RG018i004p00813 ADSGoogle Scholar
  187. A. Seiff, D.B. Kirk, Structure of the atmosphere of Mars in summer in mid-latitudes. J. Geophys. Res. 82, 4364–4378 (1977) ADSGoogle Scholar
  188. S.P. Seth, V. Brahmananda Rao, C.M. Esprito Santo, S.A. Haider, V.R. Choksi, Zonal variations of peak ionization rates in upper atmosphere of Mars at high latitude using Mars Global Surveyor accelerometer data. J. Geophys. Res. 11, A09308 (2006a) ADSGoogle Scholar
  189. S.P. Seth, U.B. Jayanthi, S.A. Haider, Estimation of peak electron density in the upper ionosphere of Mars at high latitude (50°-70°N) using MGS ACC data. Geophys. Res. Lett. 33, L19204 (2006b) ADSGoogle Scholar
  190. V. Sheel, S.A. Haider, Calculated production and loss rates of ions due to impact of galactic cosmic rays in the lower atmosphere of Mars. Planet. Space Sci. 63/64, 94–104 (2012) ADSGoogle Scholar
  191. V. Sheel, S.A. Haider, P. Withers, K. Kozarev, I. Jun, S. Kang, G. Gronoff, C. Simon Wedlund, Numerical simulation of the effects of a solar energetic particle event on the ionosphere of Mars. J. Geophys. Res. 117, A05312 (2012) ADSGoogle Scholar
  192. H. Shinagawa, S.W. Bougher, A two-dimensional MHD model of the solar wind interaction with Mars. Earth Planets Space 51, 55–62 (1999) ADSGoogle Scholar
  193. H. Shinagawa, T.E. Cravens, A one-dimensional multispecies magnetohydrodynamic model of the day side ionosphere of Mars. J. Geophys. Res. 94, 6506–6516 (1989) ADSGoogle Scholar
  194. M.D. Smith, Interannual variability in TES atmospheric observations of Mars during 1999–2003. Icarus 167, 148–165 (2004) ADSGoogle Scholar
  195. M.D. Smith, THEMIS observations of Mars aerosol optical depth from 2002–2008. Icarus 202, 444–452 (2009) ADSGoogle Scholar
  196. M.D. Smith, J.C. Pearl, B.J. Conrath, P.R. Christensen, Thermal Emission Spectrometer results: Mars atmospheric thermal structure and aerosol distribution. J. Geophys. Res. 106, 945–956 (2001) Google Scholar
  197. M.D. Smith, G. Neumann, R.E. Arvidson, E.A. Guinness, S. Slavney, Mars global surveyor laser altimeter mission experiment gridded data record IEG025_A.TAB, MGS-M-MOLA-5-MEGDR-L3-V1.0, NASA Planetary Data System (2003) Google Scholar
  198. H. Spinrad, G. Munch, L.D. Kaplan, The detection of water vapor on Mars. Astrophys. J. 137, 1319–1321 (1963) ADSGoogle Scholar
  199. A.L. Sprague et al., Interannual similarity and variation in seasonal circulation of Mars’atmospheric Ar as seen by Gamma Ray Spectrometer on Mars Odyssey. J. Geophys. Res. 117, E04005 (2012). doi:10.1029/2011JE003873 ADSGoogle Scholar
  200. S. Tellmann, M. Patzold, B. Hausler, D.P. Hinson, G.L. Tyler, The structure of Marslower atmosphere from Mars Express Radio Science (MaRS) occultation measurements. J. Geophys. Res. 118, 306–320 (2013). doi:10.1029/jgre.20058 Google Scholar
  201. W.K. Tobiska, T. Woods, F. Eparvier, R. Viereck, L. Floyd, D. Bouwer, G. Rottman, O.R. White, The solar 2000 empirical solar irradiance model and forecast tool. J. Atmos. Sol.-Terr. Phys. 62, 1233–1250 (2000). doi:10.1016/S1364-6826(00)00070-5 ADSGoogle Scholar
  202. A.H. Treiman, J.S. Treiman, Cometary dust streams at Mars: Preliminary predictions from meteor streams at Earth and from periodic comets. J. Geophys. Res. 105, 24571–24581 (2000) ADSGoogle Scholar
  203. G.L. Tyler et al., MGS RST science data products, MGS-M-RSS-5-SDP-V1.0, vol. USA_NASA_JPL_ MORS_1038, NASA Planetary Data System (NASA Goddard Space Flight Center, Greenbelt, 2007) Google Scholar
  204. A. Valeille, M.R. Combi, S.W. Bougher, V. Tenishev, A.F. Nagy, Three-dimensionalstudy of Mars upper thermosphere/ionosphere and hot oxygen corona: 2 solar cycle, seasonal variations, and evolution over history. J. Geophys. Res. 114, E11006 (2009a). doi:10.1029/2009JE003389 ADSGoogle Scholar
  205. A. Valeille, V. Tenishev, S.W. Bougher, M.R. Combi, A.F. Nagy, Three-dimensional study of Mars upper thermosphere/ionosphere and hot oxygen corona. J. Geophys. Res. 114, E11005 (2009b). doi:10.1029/2009JE003388 ADSGoogle Scholar
  206. M.B. Vasiliev et al., Preliminary results of dual frequency radio occultation of the Martian ionosphere with the aid of Mars 5 spacecraft. Kosm. Issled 13, 48–51 (1975) ADSGoogle Scholar
  207. M.I. Verigin, K.I. Gringauz, N.M. Shutte, S.A. Haider, K. Szego, P. Kiraly, A.F. Nagy, T.I. Gombosi, On the possible source of the ionization in the nighttime Martian ionosphere 1. Phobos 2 HARP electron spectrometer measurements. J. Geophys. Res. 96, 19307–19313 (1991) ADSGoogle Scholar
  208. B.P. Weiss, L.E. Fong, H. Vali, E.A. Lima, F.J. Baudenbacher, Paleointensity of the ancient Martian magnetic field. Geophys. Res. Lett. 35, L23207 (2008). doi:10.1029/2008GL035585 ADSGoogle Scholar
  209. R.C. Whitten, L. Colin, Ionosphere of Mars and Venus. Rev. Geophys. 12, 155–192 (1974) ADSGoogle Scholar
  210. R.C. Whitten, I.G. Poppoff, J.S. Sims, The ionosphere of Mars below 80 km altitude—I. Quiescent conditions. Planet. Space Sci. 19, 243–250 (1971) ADSGoogle Scholar
  211. R.C. Whitten, W.J. Borucki, S.N. Tripathi, Predictions of the electrical conductivity and charging of the aerosols in Titan’s nighttime atmosphere. J. Geophys. Res. 112, E04001 (2007). doi:10.1029/2006JE002788 ADSGoogle Scholar
  212. P. Withers, A review of observed variability in the dayside ionosphere of Mars. Adv. Space Res. 44, 277–307 (2009) ADSGoogle Scholar
  213. P. Withers, Attenuation of radio signals by the ionosphere of Mars: Theoretical development and application to MARSIS observations. Radio Sci. 46, RS2004 (2011) ADSGoogle Scholar
  214. P. Withers, M. Mendillo, Response of peak electron densities in the Martian ionosphere to day-to-day changes in solar flux due to solar rotation. Planet. Space Sci. 53, 1401–1418 (2005) ADSGoogle Scholar
  215. P. Withers, M.D. Smith, Atmospheric entry profiles from the Mars exploration rovers spirit and opportunity. Icarus 185, 133–142 (2006) ADSGoogle Scholar
  216. P. Withers, S.W. Bougher, G.M. Keating, The effects of topographically-controlled thermal tides in the martian upper atmosphere as seen by the MGS accelerometer. Icarus 164, 14–32 (2003) ADSGoogle Scholar
  217. P. Withers, M. Mendillo, D.P. Hinson, K. Cahoy, Physical characteristics and occurrence rates of meteoric plasma layers detected in the Martian ionosphere by the Mars global surveyor radio science experiment. J. Geophys. Res. 113, A12314 (2008). doi:10.1029/2008JA013636 ADSGoogle Scholar
  218. P. Withers, M.O. Fillingim, R.J. Lillis, B. Häusler, D.P. Hinson, G.L. Tyler, M. Pätzold, K. Peter, S. Tellmann, O. Witasse, Observations of the nightside ionosphere of Mars by the Mars Express Radio Science Experiment (MaRS). J. Geophys. Res. 117, A12307 (2012) ADSGoogle Scholar
  219. Ma. Yueha, I.P. Williams, W.H. Ip, W. Chen, The velocity distribution of periodic comets and the meteor shower on Mars. Astron. Astrophys. 394, 311–316 (2002) ADSGoogle Scholar
  220. M.H.G. Zhang, J.G. Luhmann, A.J. Kliore, An observational study of the nightside ionosphere of Mars and Venus with radio occultation methods. J. Geophys. Res. 95, 17095–17107 (1990) ADSGoogle Scholar
  221. H. Zou, J.S. Wang, E. Nielsen, Reevaluating the relationship between the Martian ionosphere peak density and the solar radiation. J. Geophys. Res. 111, A07305 (2006) ADSGoogle Scholar

Copyright information

© Springer Science+Business Media Dordrecht 2014

Authors and Affiliations

  1. 1.Physical Research LaboratorySpace and Atmospheric Sciences DivisionAhmedabadIndia
  2. 2.CSIR, National Physical LaboratoryRadio and Atmospheric Sciences DivisionNew DelhiIndia

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