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Atmospheres and ionospheres

Neutral atmospheres

Both the Earth and Mars exert gravitational forces great enough to maintain an atmosphere. While their atmospheres di er by their composition, their pressure levels and their temperature pro les, they are essentially governed by the same equations (continuity, momentum, di usion. . . ). Physical parameters, such as temperature, pressure and density of a given species, exhibit spatiotemporal variations resulting from the boundary conditions and the external sources of energy applied to the atmosphere, and their time evolution. The main external source providing energy to planetary atmospheres is the Sun, whose radiation heats the atmospheric constituents. An immediate consequence is that typical atmospheric pro les are di erent on the dayside and nightside of a planet, and also vary according to solar activity. Other sources of energy include for instance galactic cosmic rays [e.g., Velinov et al., 2013], particle precipitation (see section 1.4.3), Joule heating, atmospheric gravity waves [Hickey et al., 2011], or volcanism [Vyushin et al., 2004].
Figure 1.1 gives typical pro les of the main atmospheric parameters for the Earth, given by the MSIS-90 model [Hedin, 1991] above Sodankyl on 1 January 2014 at 12:00 local time. Focusing rst on the temperature pro le, one may notice several vertical gradient inversions below 100 km altitude, which de ne four distinct regions, called with increasing altitude: troposphere, stratosphere, mesosphere, and thermosphere. Above 100 km altitude, i.e., in the thermosphere, the in uence of the 11-year solar cycle is signi cant. The exospheric temperature (i.e., the temperature in the outermost regions of the atmosphere) may vary from about 1000 K during so-called solar minimum conditions to 1400 K during so-called solar maximum conditions [Brekke, 2012]. The composition of the atmosphere varies with altitude. Below 100 km, the atmosphere is a well-mixed gas made of, roughly, 78% of nitrogen (N2), 21% of oxygen (O2) and 1% of argon (Ar) and other gases. Above 100 km, as density decreases, collisions become less frequent and each species therefore behaves independently. While the heaviest species are more a ected by gravitation and tend to remain at lower altitude, the lightest ones are prone to escaping. Atomic oxygen, formed by dissociation of O2 by solar ultraviolet (UV) radiation, appears above 90 km altitude, and becomes the main species above 150 km [Rees, 1989]. Other atomic species (He and H) become dominant in the topside atmosphere (above 600 km) [Hargreaves, 1995].
Another key parameter in a planetary atmosphere is the scale height, noted H and given by with kB the Boltzmann constant, T the temperature, M the molecular mass, and g the standard acceleration due to gravity. Under the hydrostatic equilibrium assumption, the neutral density N at altitude z can be calculated as a function of temperature by where N0, T0 are neutral density and temperature values at a reference altitude z0. In the case of an isothermal atmosphere, this equation simpli es  and the scale height therefore gives the altitude increase corresponding to a density decrease by a factor of e. Above 100 km altitude, each neutral constituent has its own scale height. The scale heights are of the order of a few kilometres below 100 km altitude, and increase to several tens of kilometres above, with a dependency on solar activity since they depend on temperature. Correspondingly, the neutral density in the thermosphere and exosphere also depends on solar activity.
At Mars, the atmosphere is governed by the same concepts, with some di erences. First, the composition of the Martian atmosphere is very di erent. Below 120 km altitude, where the atmosphere is a well-mixed gas [Izakov, 2007], the main constituent is carbon dioxyde (CO2, 96%), the remainder consisting mostly of argon and molecular nitrogen [Maha y et al., 2013]. Above 200 km altitude, atomic oxygen (O) becomes the main neutral species [Fox , 2009]. The temperature essentially decreases from ground level to about 40{120 km altitude, where it reaches a minimum [Nier and McElroy, 1977; Sei and Kirk, 1977]. Above that, the thermospheric pro le is exponential, the exospheric temperature ranging between 200 and 300 K, depending on solar activity [Fox , 2009]. Finally, the Martian atmosphere is very rare ed, with ground pressure levels of the order of 100 times lower than on Earth. Figure 1.2 gives an example of atmospheric temperature and neutral density pro les (for CO2, O and N2) obtained from the Mars Climate Database v5.2 [Millour et al., 2015] during the conditions of the landing of the Curiosity rover, on 6 August 2012.
One speci city of the Martian atmosphere is the dust which is lifted from the ground up to 60 km altitude by strong winds during the dust storm season [Clancy et al., 2010]. This dust strongly a ects the temperature and densities in the lower atmosphere, especially by absorbing solar radiation and reemitting energy in the infrared frequency range, hence heating the neutral gas.
Figure 1.2. Altitude pro les in the Martian atmosphere at the landing time and location of the Curiosity rover, given by the Mars Climate Database v5.2. Red: Neutral temperature. Blue: Number densities of CO2 (solid), O (dashed) and N2 (dash-dotted). Black: Total neutral density.
A comparison of the terrestrial and Martian neutral atmospheres shows that, besides very di erent compositions, the two main di erences noticeable in the presented parameters are that (1) pressure and total neutral density are about two orders of magnitude smaller at Mars than at Earth, and (2) the Martian atmosphere is overall colder than the terrestrial atmosphere. At ground level, the temperature di erence is only of the order of a few tens of Kelvin, but the exospheric temperature is signi cantly higher at Earth (1000{1400 K) than at Mars (200{300 K).


Photoionisation, chemistry and transport

Besides heating the atmosphere, the solar radiation also modi es its composition. Solar photons in the extreme ultraviolet (EUV) and X-ray energy ranges indeed carry enough energy to dissociate molecules and ionise constituents. As a result, a great number of photochemical and chemical reactions take place in the atmospheres of the Earth and Mars, including photodissociation, photoionisation, charge exchange and recombination reactions.


In the upper atmosphere, ion species and free electrons are always present, and this region is called the ionosphere. The ionosphere is a plasma of low ionised fraction (10 3 for the Earth, 10 5 for Mars). Ionospheres are commonly divided into regions, based on their electron density altitude pro les.
In an ionosphere, neutrals, electrons and ions behave di erently. The motion of neutral species (neutral winds) is mostly driven by pressure gradient forces. The motion of electrons mainly depends on electromagnetic elds, as their collision frequency with other species is very low. They follow the E B drift (see section 1.4.1), at a speed vE B = E B=B2 (with E the electric eld and B the magnetic eld). The motion of ion species, on the other hand, is in uenced by both electromagnetic elds and collisions with neutral particles. The ion-neutral collision frequency varies with altitude. In the lower ionosphere (up to about 200 km altitude for the Earth and Mars [Vogt et al., 2016]), collisions with the neutrals are frequent, and as a result the ions drift at a speed with a component parallel to the electric eld and a component in the E B direction. Since the ion motion is not in the same direction as the motion of the electrons, currents are owing in the plasma. At higher altitude, however, the collisions are rare enough to allow the ion motion to follow the E B drift, as for the electrons [Brekke, 2012]. This produces plasma convection [Hargreaves, 1995]. The eld-aligned motion of ions is dominated by the eld-aligned component of the neutral wind. Ambipolar di usion of electrons and ions, essentially upwards, also plays a role in the eld-aligned motion of charged particles [Rees, 1989].

Structures of the terrestrial and Martian ionospheres

The terrestrial ionosphere is mainly divided into three regions according to the electron density pro le. Figure 1.3 shows typical electron density pro les in the ionosphere of the Earth during daytime and nighttime, as well as density pro les of the main ions during daytime. Between about 60 and 90 km altitude, the so-called D region has a complex photochemistry, which includes negative and cluster ions [Turunen et al., 1996]. It is formed principally by solar Lyman- , EUV and hard X-ray radiation, galactic cosmic rays, and energetic to relativistic particle precipitation from the magnetosphere (see section 1.4.3). The D region generally disappears at night, as recombination takes place rapidly in this region where collision frequencies are high.
Between 90 and 150 km altitude, approximately, the E region predominantly consists of O+2 and NO+ ions. In this region, the motions of ions and electrons are decoupled, which allows the existence of electric currents owing in this part of the ionosphere. The E region electron density is greatly reduced at night, but at high latitudes it may also be signi cantly enhanced by auroral precipitation.
Above 150 km altitude, the F region contains the main electron density peak, which may be of the order of 1012 m 3. During solar maximum and on the dayside, the F region may exhibit a bulge under the main peak; it is then commonly divided into F1 and F2 regions. The F1 region consists mostly of molecular ions (O+2 and NO+) and disappears at night, while the dominant species in the F2 peak is O+. The main peak altitude typically ranges between 250 and 300 km during daytime, and between 300 and 400 km during nighttime [Rishbeth et al., 2000a].
Figure 1.3. Typical electron density pro les in the daytime (black, solid) and nighttime (black, dashed) terrestrial ionosphere (after Cravens [1997]). Density pro les of the main ion species in the daytime ionosphere are given in colours. The ion composition was estimated using the TRANSCAR model [Blelly et al., 2005].
In spite of the background atmosphere being di erent, the Martian ionosphere presents a structure with some similarities to the terrestrial ionosphere. The same general formation mechanisms { photoionisation, chemistry, ion and electron di usion { lead to the creation of an ionosphere with two main regions. While there is no commonly used o cial nomenclature for the Martian ionospheric regions, it is chosen here to follow Rishbeth and Mendillo [2004] and call them M1 and M2. Figure 1.4 shows a typical electron density pro le from 60 to 240 km altitude in the dayside Martian ionosphere. The main peak M2 is located at about 120 km altitude and its electron density is of the order of 1011 m 3. The secondary peak, M1, lies at 100{ 110 km altitude and resembles more a bulge than an independent peak. In fact, the M1 and M2 regions are believed to be distinct when the solar zenith angle is high and to merge to form a single electron density peak near local noon [Mayyasi and Mendillo, 2015]. The main ion species in both regions is O+2, but NO+ may become dominant below 100 km altitude [Fox , 2009]. The nightside ionosphere of Mars is signi cantly less well-known, but recent studies suggest that the ionisation does not completely disappear at night [Withers et al., 2012; Fowler et al., 2015]. Besides, it has been found that the global dust storms, by a ecting the neutral atmosphere (see section 1.1.1), may also have an e ect on the ionosphere. Indeed, the dust-loaded In addition to the main ionospheric region consisting of M1 and M2, some studies have revealed a sporadic layer at lower altitude, supposedly originating from meteoric ablation and therefore consisting of metallic ions (mostly Mg+ and Fe+) [Molina-Cuberos et al., 2003]. This layer is located around 80 km altitude, and it almost disappears on the nightside. Models have also predicted the existence of ionisation associated with cosmic rays down to about 35 km altitude [Withers, 2009; Cardnell et al., 2016].
The main di erences between the terrestrial and Martian ionospheres can be summarised as follows. First, the altitude extent of the ionosphere of Mars is reduced compared to the terrestrial case, as the main peak is located around 120 km (versus 250{300 km at Earth) during daytime. Second, the main peak electron density is one order of magnitude lower at Mars than at Earth ( 1011 vs 1012 m 3). Third, the main ion species at most altitudes in the Martian ionosphere is O+2, whereas in the ionosphere of the Earth O+ dominates in the main peak and above, and NO+ dominates in the E and lower F regions.
These descriptions of the terrestrial and Martian ionospheres correspond to typical, undisturbed conditions, and the presented electron density pro les are idealised ones. However, ionospheres are subject to great dynamical variations triggered by disturbances in the nearby space. In the case of the Earth, for instance, there is a strong coupling between the ionosphere, the magnetosphere (described in section 1.3.1) and the solar wind. At Mars, a coupling between the ionosphere and the solar wind also exists, and will be described in section 1.4.4.

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Solar wind

General properties

The Sun also a ects the planetary environments through the continuous expansion of its own atmosphere, the corona. This plasma ow is called the solar wind, and it is the main driver of space weather.
The outward ow of the solar wind originates from the pressure gradient existing between the solar surface and the outer solar atmosphere, whose e ect is greater than the gravitational forces exerted on the coronal constituents. The main constituents of the solar wind are protons and electrons, and about 5% of the positive charge careers are -particles (He2+) [Hundhausen, 1995]. When measured near the Earth’s orbit, the average solar wind ion density is of the order of 3{10 particles per cubic centimetre. Its velocity is predominantly radial and typically ranges between 200 and 800 km/s [Hargreaves, 1995]. The dynamic pressure, given by pdyn = V 2 (with the mass density and V the velocity) [Walker and Russel, 1995], is therefore most of the time of the order of a few nanopascals. At Mars, the Mars Atmosphere and Volatile Evolution (MAVEN) mission recently started to provide solar wind observations using the Suprathermal and Thermal Ion Composition instrument. A statistical study by Masunaga et al. [2017] using 1.3 year of measurements estimated that ion density, solar wind velocity and dynamic pressure have values at Mars of the order of 1{10 cm 3, 200{600 km/s, and 0.2{2.0 nPa, respectively.
The solar wind being a plasma originating from the Sun, which has an intrinsic magnetic eld, it drags the solar magnetic eld along its outward motion. This comes from the fact that the solar wind conductivity along the magnetic eld lines is very high, and this is called the \frozen-in » ux concept [e.g., Baumjohann and Treumann, 1997]. The magnetic eld carried by the solar wind is called interplanetary magnetic eld (IMF). Since the Sun rotates around its spin axis and the solar wind mostly ows radially, the interplanetary magnetic eld lines draw a spiral in the interplanetary space, which is called the Parker spiral, after Parker [1958]. At the Earth and Mars orbits, the spiral angles made with the Sun{planet direction are close to 45 and 57 , respectively. The IMF strength is typically of the order of 5 nT near the Earth [Brekke, 2012] and 3 nT near Mars [Masunaga et al., 2017], yet values may commonly range from 1 to 15 nT.
The solar wind exhibits great spatiotemporal variability, both during quiet and disturbed conditions. Disturbances in the background plasma ow may be created by structures forming on the solar surface. The most extreme type of disturbances arises when solar plasma is released from the surface of the Sun, forming a so-called coronal mass ejection (CME). Such a structure exhibits a large magnetic eld magnitude and often travels at high speed (exceptionally up to 2500 km/s) in the interplanetary medium [Lugaz et al., 2005]. CMEs are most frequent during the period of maximum activity in the 11-year solar cycle, and they are often released following a sudden brightening of a localised region at the surface of the Sun, called a solar are [Hargreaves, 1995].

High-speed streams

Another type of solar wind structures a ecting space weather, the solar high-speed streams are created by the fast ow of solar wind plasma along open magnetic eld lines. These regions of open magnetic ux are called coronal holes, and they are almost always present on the Sun, especially at high latitudes [Lowder et al., 2017]. During the declining phase of the solar cycle, coronal holes from the polar regions tend to migrate to lower latitudes, and the resulting high-speed streams of solar wind may ow in the ecliptic plane [Kojima and Kakinuma, 1987]. When the fast solar wind (typically 500{800 km/s) is preceded by a slower stream (200{400 km/s), this creates a compression region characterised by high IMF magnitude, solar wind density and pressure, called corotating interaction region [Richter and Luttrell, 1986]. The terminology came from the observation that such structures often show a 27 day recurrence, corresponding to the solar rotation period near the equator, which implies that a same coronal hole (and hence, the resulting high-speed stream) may persist during several solar rotations. Figure 1.5 shows the IMF magnitude and the solar wind velocity, density and dynamic pressure measured near the Earth during a high-speed stream event from March 2007. The transition from the slow solar wind stream to the high-speed stream contains several local peaks of solar wind pressure and density, in this particular example. The east-west component of the solar wind velocity (not shown here) generally exhibits a sign reversal at the transition from slow to fast solar wind, de ning the stream interface [Kavanagh et al., 2012].

Table of contents :

1 Introduction 
1.1 Atmospheres and ionospheres
1.1.1 Neutral atmospheres
1.1.2 Ionospheres Photoionisation, chemistry and transport Structures of the terrestrial and Martian ionospheres
1.2 Solar wind
1.2.1 General properties
1.2.2 High-speed streams
1.3 Magnetospheres
1.3.1 Terrestrial magnetosphere
1.3.2 Martian induced magnetosphere
1.4 Planetary environments and solar wind interactions
1.4.1 Reconnection and convection
1.4.2 Geomagnetic storms and magnetospheric substorms
1.4.3 Particle precipitation
1.4.4 Some aspects of Mars interactions with the solar wind
2 Instrumentation 
2.1 Ground-based instruments
2.1.1 Sodankyla ionosonde
2.1.2 IMAGE magnetometers
2.1.3 SGO riometer chain and KAIRA
2.1.4 Kilpisjarvi All-Sky Camera
2.2 Satellites
2.2.1 Mars Express
2.2.3 ACE
3 Methods 
3.1 Superposed epoch analysis
3.1.1 Classical version
3.1.2 Phase-locked version
3.2 Radio-occultation
3.2.1 Principle
3.2.2 Data analysis
4 Results 
4.1 Efects of solar wind high-speed streams on the high-latitude ionosphere
of the Earth
4.1.1 Solar wind driving
4.1.2 Overview of ionospheric response to high-speed streams
4.1.3 E and F region responses
4.1.4 Energetic particle precipitation
4.2 Modulation of energetic precipitation during pulsating aurora
4.3 Analysis of Mars Express data with the radio-occultation model
4.3.1 Reproduction of the frequency residual prole
4.3.2 Inuence of medium asymmetry
4.3.3 Ion density proles
4.3.4 Improvements introduced in Paper
5 Conclusion 


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