Chapter I. Introduction to Atmosphere and Ionosphere

Transcription

1 Chapter I Introduction to Atmosphere and Ionosphere 1.1. Introduction The gaseous envelope surrounding the earth is its atmosphere whose uppermost region is weakly ionized by the high energy electromagnetic radiation coming from the sun. The region where the ionization due to solar radiation occurs is called the ionosphere. The important physical properties that determine the state of the atmosphere are temperature, pressure, density and wind, while for the ionosphere plasma density is also important in addition to the abovementioned properties. The atmospheric density and pressure decreases exponentially with height. As a result, an upward directed pressure gradient is generated and is balanced by the downward directed gravitational force. This balance between the pressure gradient force and gravity is called Hydrostatic balance. Equation for hydrostatic balance is, P z = ρg...(1.1) where, P and represent pressure and density respectively, g is the acceleration due to gravity and z represents the altitude. Another useful quantity is the scale height, H, defined as, kt H =...(1.) mg where, k is the Boltzman constant (1.38x10-3 J/K), T is the temperature and m is the weighted molecular mass of the atmospheric constituents at the respective height. Scale height represents the height over which atmospheric pressure (and density) falls by a factor of e. Similar to the neutral species, plasma scale height is defined in the ionosphere with 1

2 temperature (T) in eqn. (1.) replaced by sum of electron and ion temperatures (T e +T i ) and mass representing the weighted molecular mass of positive ions. Potential temperature is a useful concept in studying the atmospheric dynamics and energetics. It is defined as the temperature of an air parcel when it is adiabatically compressed to a pressure of mb. It is given by, θ P T P = 0 R C P...(1.3) where, is the potential temperature, P 0 is the reference pressure of 1000 mb and P is the pressure at the particular altitude of air parcel (in millibars), R is the universal gas constant and C p is the specific heat at constant pressure. The potential temperature of an air parcel is a conserved quantity for an adiabatic process. Fig The composition of major atmospheric species (adapted from Hargreaves, 199).

3 The composition of atmosphere is fairly constant up to about 100 km with approximately 78% of nitrogen, 1% of oxygen and 0.9% of argon. In the lowermost regions of atmosphere, the concentration of water vapour is highly variable and may account for about 1% in the regions of intense convection. All other trace constituents comprise the remaining meagre part. Many of the trace gases are chemically active, thereby affecting the energetics of the atmosphere. The composition of the atmosphere neglecting the trace species is given in Figure 1.1. From the dynamical stand point, the atmosphere can be considered as an ideal gas to a high degree of approximation. The equations derived by considering the atmosphere as an ideal gas elegantly explain the dynamics of the atmosphere. The region below 100 km altitude is dominated by waves and instabilities which bring about turbulence and mixing thereby maintaining the uniform composition of major species in the atmosphere. This region is called homosphere. The region above homosphere is dominated by the diffusive separation of gases based on their molecular mass. The diffusive and dissipative effects suppress dynamical activity in this region which results in insignificant mixing of the constituents. This region is called heterosphere. The boundary between these two regions is called homopause or turbopause. Apart from this broad classification, the atmosphere can also be classified into different layers based on its temperature structure as described in section 1.. Ionosphere starts at altitudes of ~60 km. The ionosphere is separated into three distinct layers based on the electron density variations and the composition of ionic species. Different kinds of phenomena occur in each layer. The classification of ionosphere is briefly discussed in section 1.3. Section 1.4 provides an introduction to the airglow and discusses prominent nightglow emissions that are observed from the ground. Atmospheric gravity waves and the 3

4 plasma irregularities resulting from Rayleigh-Taylor instabilities occurring in the night time ionosphere are concisely described in sections 1.5 and 1.6, respectively. Fig.1.. The temperature structure of the atmosphere 1.. Thermal structure of the atmosphere Figure 1. shows the temperature structure of the atmosphere. The alternate regions of negative and positive temperature gradients are classified as different regions as discussed below. The layers are called spheres and their boundaries are called pauses. It may be noted that the boundary between the regions are not well defined and they may vary by few kilometres in both space and time. 4

5 1..1. Troposphere and stratosphere The lowest and the densest part of the atmosphere is the troposphere. It extends to about 8 km over the poles and 16 km over the equator. The troposphere contains roughly 85% of the mass of atmosphere. In troposphere the temperature decreases upward. The heat source of this layer is the reradiated long wave IR radiation (in the wavelength region of m) from the surface of the earth. The temperature decreases with height as a result of reduction in the intensity of reradiated IR radiation and decreasing density of air. The decrease in temperature with height is termed as lapse rate. For dry air, the lapse rate is called dry adiabatic lapse rate, = g/c p. The value of dry adiabatic lapse rate is approximately 9.5 K/km. The lapse rate is considerably different for moist air due to the heat content of water vapour. A substantial portion of water vapour in the atmosphere is present in the troposphere. The environmental lapse rate in the troposphere is 6 7 K/km. The decreasing temperatures eventually reach a minimum of ~10 K in the top of the troposphere after which a temperature increase is noticed. This region of temperature minimum is called tropopause. The region above the tropopause up to about 50 km is marked by increase in temperature with height and is called stratosphere. Since temperature rises with height convective activity is effectively prohibited resulting in considerably reduced vertical mixing and turbulence. The increase in temperature results from the absorption of solar ultraviolet radiation between 10 and 310 nm by dissociation of ozone molecules (O 3 ). Even though the ozone concentrations are largest in the 0-30 km region, the temperature maximum occurs around 50 km region as a result of balance between the amount of influx of UV rays and ozone concentration. The top of the stratosphere is called the stratopause above which the temperature decreases with height. 5

6 1... Mesosphere and thermosphere Mesosphere is the region of the atmosphere above stratopause where the temperature decreases again, similar to the troposphere. This layer extends up to about 100 km except over the summer high latitude regions. The temperature reaches a minimum at the top of the mesosphere in the region called mesopause, which is the coldest region in the entire atmosphere. The temperatures fall below 180 K in the mesopause region. In summer high latitude regions, the mesopause is found at lower altitudes of ~87 km with further reduced temperatures, typically below 150 K. This latitudinal variation in the altitude of mesopause is referred to as two-level mesopause. A particular type of clouds called polar mesospheric clouds (PMCs) or noctilucent clouds form in the vicinity of summer polar mesopause owing to the extreme low temperatures of the region. There is no effective solar absorption in the mesosphere to act as a heat source. Some insignificant heat contribution in this region comes from absorption in the Schumann-Runge bands at wavelength between 170 and 195 nm by O. Apart from the fact that this region does not have significant heat source, considerable energy is lost by means of radiative cooling, especially due to the trace gases carbon dioxide, nitric oxide and ozone. The mesospheric region is dynamically very active. Internal gravity waves, atmospheric tides and planetary waves are observed in this region and they arrive from lower atmosphere. As a result of conservation of kinetic energy, gravity waves and tides attain large amplitudes when they reach mesospheric altitudes of lower atmospheric density. The amplitude of internal gravity waves becomes large enough to undergo saturation and breaking. When the waves break, they deposit their energy and momentum in those regions. Further, the wave breaking and wave-wave interactions involving internal gravity waves are found to trigger instabilities in the mesosphere (discussed in Chapter IV), which lead to 6

7 turbulence and mixing. The internal gravity waves are the most important dynamical features that drive the mesospheric circulation. Indeed, the latitudinal temperature structure of upper mesosphere is contradictory to the radiative equilibrium in that the winter polar regions are warmer than the summer polar regions. This discrepancy is attributed to the dynamical circulation of mesosphere induced by internal gravity waves. In addition, the dynamics of the upper mesospheric region appear to play a fundamental role in the variabilities occurring in the ionospheric E region (to be discussed in the next section (1.3)). More than 100 metric tons of meteoric debris enter the atmosphere per day, most of them in the form of small meteoroid particles with sizes smaller than 1 mm. They contain metal traces and their evaporation occurs at the upper mesospheric heights resulting in formation of atomic metallic layers between 80 and 100 km. Apart from these, a part of the airglow and auroral emissions emanate from the upper mesospheric region. This thesis extensively discusses various types of small-scale dynamical events that occur in the upper mesospheric heights. Such features are believed to play dominant role in determining the local variations in the upper mesosphere. Other fascinating features occurring in this region are lightning induced transient luminous events like red sprites and ELVEs. The region above the mesosphere where the temperature increases sharply and becomes constant at about km is called the thermosphere. The absorption of shorter wavelength solar photons in extreme ultraviolet (EUV) and X-ray portions of the electromagnetic spectrum by the tenuous gas in the thermosphere is the primary heat source of this region. At high latitudes, Joule heating associated with the electric currents driven by magnetospheric electric fields also contributes to heating. At altitudes of lower thermosphere, 7

8 below about 170 km, the photodissociation of O by the Schumann-Runge continuum ( nm) is an important heat source. Most of the heat liberated in the upper thermosphere is removed by downward conduction which results in a drastic rise of temperature in the lower thermosphere region. The heat conductivity becomes too good in the upper thermosphere that it is maintained in nearly isothermal condition at a relatively high temperature of several 100s K. thermospheric temperatures vary with solar activity as the irradiance of high energy photons also vary with solar activity. Many airglow and auroral emissions emanate from the thermosphere. In addition to the downward heat conduction, radiative cooling by airglow emissions and nitric oxide are also other energy loss mechanisms. Uneven heating over different regions of thermosphere generates pressure gradients that give rise to neutral wind motions which tend to redistribute energy. The region above the thermosphere is termed as exosphere. It is not considered to be a part of the atmosphere, but the base of the exosphere called exobase or baropause is traditionally considered as the boundary of the atmosphere.. During solar maximum periods the exospheric temperature can rise to 1800 K while it will be ~800 K during solar minimum conditions. The notion of fluid is not applicable to the tenuous gas constituting the exosphere because the mean free paths are of the order of scale heights in this region. Neutral atoms move in ballistic orbits entirely under the influence of gravity whereas ionized particles are constrained by the geomagnetic field. 8

9 Fig Composition of ions and neutral atoms in ionospheric region (adapted from Kelley, 009) 1.3. Ionosphere As already mentioned, the weakly ionized part of the upper atmosphere constitutes the ionosphere whose plasma density never exceeds 1% of neutral density at any heights. Figure 1.3 shows the composition of the important ions and neutrals at ionospheric altitudes. The ionosphere coexists with mesosphere and thermosphere regions. Though the ionosphere is weakly ionized, the electrodynamical aspects associated with electrons and positive ions gain importance over the neutral dynamics in the region. The ionosphere is permeated by the geomagnetic field. The combined action of neutral winds and magnetic field generates electric fields in the ionosphere which give rise to current systems. The ionosphere is classified into three regions as discussed below briefly [Rishbet and Garriot, 1969; Kelley, 009]. Figure 1.4 shows the electron densities of different regions of the ionosphere during day and night time. 9

10 Fig Electron density profile of the ionospheric region. Solid lines represent solar maximum conditions and dashed lines represent solar minimum conditions D and E regions The D region is the innermost region occurring between 60 and 90 km above the surface of the earth. In this region, the ionization results mostly from photoionization of nitric oxide (NO) by Lyman radiation of wavelength 11.6 nm. In addition, hard x-rays of wavelength less than 1 nm ionizes N and O. The lower part of the layer below 75 km is ionized by cosmic rays. The primary ions present in this region are the molecular ions NO +, O + and N +. In addition, owing to the relatively higher neutral density compared to other regions of ionosphere, cluster ions are formed in this region. Also, some of the electrons attach to the neutrals forming negative ions. D region plasma density is in the range cm -3. The plasma in this region recombines after sunset except for small amount of ionization that remains due to galactic cosmic rays. In the D region, the electron-neutral collision frequency is about 10 6 s -1. Thus, the amplitude modulated radio wave transmissions are highly absorbed during the day and almost unattenuated at night. This region is dominated by 10

11 neutral dynamics due to frequent collisions and electrodynamics does not play any important role. The E region was the first identified electrically conducting region in the upper atmosphere. It extends from ~90 to ~150 km. The source of the E region ionization is soft X- rays of wavelength between 1-10 nm and far ultraviolet (1 00 nm) solar radiation that ionizes molecular oxygen. The primary ions are N +, O +, NO + and O +. Atomic metal ions like Fe +, Ca +, Si + and Mg + are also observed in the lowest parts of this region in meagre concentrations. These ions are produced by meteor ablation. The day time electron concentration of E region is ~10 5 cm -3 which reduces to ~10 3 cm -3 during night time. In this part of the ionosphere, the ions are still dominated by collisions with neutral species while the electrons follow geomagnetic field lines. The atmospheric neutral winds dominated by tidal modes force the ion component of the E-region to move across the geomagnetic field lines. This differential motion between ions and electrons result in electric fields. These electric fields in turn give rise to current systems like equatorial electrojet and solar quiet (S q ) current system. At high latitudes, the electric fields of magnetospheric origin map to the E region heights and generate current systems known as auroral electrojets that are enhanced significantly during magnetic storms and substorms. During night time, the electron density and conductivity of the E region decrease substantially F region The F region is the uppermost region of the ionosphere starting from about 150 km. The F region does not have a fixed upper boundary. It extends up to about 1000 km and merges with the protonosphere. This region coexists with the thermosphere and a part of it extends into the exosphere as well. Ionization in this region results from the extreme ultraviolet radiation between 10 and 11 nm which ionizes atomic oxygen and molecular 11

12 nitrogen. The principal ion in F region is O + ion. The electron density attains a maximum of about 10 6 cm -3 during day time and that at night time is about 5x10 4 cm -3. On several days, the daytime F region splits into two layers called F1 and F layers. The F layer is the densest part of the ionosphere. The region above F peak altitude is called topside ionosphere. Recent studies revealed occasional further layering in the top side ionosphere above F layer. This layer is termed as F3 layer and it appears to be related to intense ExB drifts occurring over the magnetic equator [Balan and Bailey, 1995]. F region typically affects propagation of radio waves in the frequency range to 16 MHz. However, this region is susceptible for the growth of large scale instabilities during night time. During such periods, depending on the evolution of instabilities, radio frequencies up to few GHz may get affected. Since this thesis concentrates on the processes occurring in the low and equatorial latitudes, we briefly describe the plasma drift and neutral wind pattern in equatorial latitudes. A detailed description of the causes and effects of the observed patterns could be found in Kelley [009] and Herrero et al. [1993]. Figure1.5 shows the F region zonal plasma drift velocities along with the thermospheric zonal wind speeds obtained by Jicamarca radar and Dynamics Explorer satellite, respectively. The general agreement between the directionalities of plasma drifts and neutral wind may be noted. However, the neutral wind velocities are relatively higher than plasma drift velocities. The directionality of neutral wind can easily be explained by the pressure gradient force that exists between day and night time thermosphere. The daytime westward and nighttime eastward zonal plasma drifts are due to the upward and downward vertical electric fields generated in the F region, respectively. The F region electric fields have their origin in the UxB drift, U being the thermospheric zonal wind, and the process is referred to as F region dynamo [Kelley, 009]. However, during daytime, F region electric 1

13 fields are short circuited by E-region conductivity and therefore overshadowed by stronger E region fields. Fig Typical F-region zonal plasma drifts and thermospheric zonal winds (adapted from Kelley, 009) Figure1.6 shows the vertical plasma drift velocities in the height range of about km over Jicamarca during different seasons and solar activity period. Upward drifts correspond to positive values in the plots. The figure shows that plasma drifts are upward during the day and downward during the night. The drifts are ExB drifts with upward drifts corresponding to the eastward zonal electric field during daytime, and vice versa. Before the reversal occurs during post sunset hours, the vertical drift attains a large enhancement known as prereversal enhancement (PRE) that correspond to a large increase in the eastward electric field. This enhancement has its origin in the interaction of thermospheric zonal wind (which is eastward in the evening) with the local time gradient in the height integrated Pedersen conductivity of the E layer that exists across the sunset terminator [Kelley, 009] ]. 13

14 Fig Vertical drift of F-region plasma measured from Jicamarca (taken from Kelley, 009). The F-region plasma is constrained to move along the geomagnetic field as the conductivity is maximum along the magnetic field lines. The field lines indeed act as equipotentials thereby mapping the E region electric fields and their variabilities to the F- region. Figure 1.7 shows the behaviour of thermospheric meridional winds in the low latitude region during solstice periods obtained from measurements of Atmosphere Explorer E satellite [Herrero et al., 1988; 1993]. The meridional winds are generally poleward during daytime and an enhancement in the poleward wind speed was found during winter period late afternoons. The wind reverses equatorward after sunset and again turns poleward around midnight in summer and after midnight in winter. The abatement of equatorward meridional 14

15 wind around midnight and its reversal to the poleward direction are found to be consistent with a phenomenon called midnight pressure bulge and midnight temperature maximum (MTM) [Colerico and Mendillo, 00]. Fig Meridional winds in low latitude region for solstice periods (adapted from Herrero et al., 1993) Another interesting low latitude nighttime ionospheric phenomenon is the equatorward movement of the equatorial ionization anomaly (EIA) arising out of reverse 15

16 fountain effect. In the daytime low latitude ionosphere, an eastward electric field lifts the plasma to sufficiently higher altitudes over the magnetic equator where recombination rates are smaller. The accumulated plasma diffuses away from the equator along the field lines in both the hemispheres due to pressure gradient and gravitational forces. This brings about a plasma enhancement typically between 15 o 0 o dip latitudes away from the dip equator. This trough of plasma density over the magnetic equator and two crests on both the sides together are referred to as the equatorial ionization anomaly (EIA). The process that brings about the EIA is referred to as plasma fountain effect. In the night time, the primary electric field turns westward after undergoing pre-reversal enhancement (PRE). The westward zonal electric field in the night time brings about a downward drift of plasma over the equator. Consequently, a plasma flow pattern that is opposite to the daytime fountain effect occurs with an equatorward movement of EIA crest region in the night time ionosphere. This is known as reverse fountain effect. Another way to conceive the forward and reverse fountain effects is by means of frozen-in-field concept of plasmas. In this view, the poleward movement and the formation of EIA can be visualized as expansion of the magnetic flux tubes while the nocturnal equatorward movement associated with reverse fountain is equivalent to the compression and squeezing in of the flux tubes at the equator. A schematic showing the formation of EIA is given in Figure 1.8. If strong thermospheric transequatorial meridional winds are present, then the electron density distribution will not be symmetric about the magnetic equator. Such a wind acts to transport plasma up the field lines in the summer hemisphere and down along the field lines in the winter hemisphere. This causes an asymmetry of equatorial anomaly peaks between the two hemispheres as shown in Fig.1.8, with a high plasma concentration on winter hemispheric crest than that in the summer hemisphere. Meridional winds directed either 16

17 poleward or equatorward on either of the hemispheres will not be effective in causing the asymmetry as it is not transequatorial in nature. Fig Schematic showing the formation of EIA crest Airglow The light sources located outside and within the atmosphere illuminate the night sky. If the light from these sources like moon, stars and lights from ground were eliminated, the sky would not be completely black. A faint glow would remain, that has its origin in atmospheric photochemical process and to which the name airglow is given. Airglow is the emission of photons from upper atmospheric constituents excited in a direct or indirect way by the electromagnetic radiation from the Sun. Examples of direct excitation are the resonance emissions of alkali metals that are observed in the twilight. 17

18 Indirect excitations are associated with chemical reactions resulting from processes like recombination of ionized or dissociated particles. Airglow emissions are a feature of most planetary atmospheres. The light in both airglow and aurora consists of atomic spectral lines and molecular spectral bands of atmospheric constituents, both neutral and ionized. Many of those spectral lines result from transitions that are forbidden by the selection rules of quantum mechanics and have long excited state life times. In spite of this general similarity, important differences exist between airglow and aurora. Aurora is energised by charged particles entering the atmosphere from the magnetosphere. Usually it occurs in high-latitude regions during and after solar disturbances. Auroras may have noticeable structural features. In contrast, airglow occurs continuously and is omnipresent in all latitudes. But it is weak and unstructured. On a moonless night the airglow contributes the major part of the light arriving from the sky. It exceeds the starlight in total intensity though its presence is not generally appreciated because of its uniform distribution across the sky. The airglow brightness is usually measured in Rayleigh units represented by the letter R. One Rayleigh is equal to 10 6 photons/cm /s 1. Airglow emissions arise when an excited atom or molecule returns to the ground state by emitting a photon. Instead, if the energy of the excited state is lost by means of collisions, it is called quenching. The most important processes that give rise to airglow are: i) Photo excitation and photoelectron excitation, ii) Radiative and dissociative recombination, iii) Collisional excitations and energy exchange and iv) Ionization excitation. Based on the time of observation, airglow is classified as below. 18

19 i) Dayglow: The airglow emissions emitted when sunlight enters the atmosphere between solar zenith angles of 0 and 90. Dayglow is dominated by direct excitation by solar photons and photoelectron impact excitations. Dayglow emission intensities are of the order of mega Rayleighs. In spite of the higher intensities, it is relatively difficult to observe dayglow emissions because of the much higher intensities of the background atmospheric scattered sunlight. ii) Twilight glow: Twilight glow occurs when the upper atmosphere is illuminated by sunlight while the ground is in darkness (i.e. sun is below the horizon). The solar zenith angles lie between 90 and 110 during twilight period. The dayglow and twilight glow excitation mechanisms are similar. However during twilight, direct solar illumination does not extend to the lower levels of the atmosphere where most of the scattering occurs and hence the level of background sky brightness is substantially reduced. Twilight glow emissions may have intensities of several kilo Rayleighs. given below The important excitation processes in the formation of dayglow and twilight glow are X + h, Resonance scattering X + h X *...(1.4) X + W + h e, Flourescence scattering In this process the excited species will return to ground state either by spontaneous radiation of the photon of same energy (resonance scattering), or go through an intermediate state by which it loses a part of the energy W (fluorescence scattering). XY + h X * + Y X + Y + h e, Photodissociation...(1.5) 19

20 X + h X +* + e + h e, Photoionization...(1.6) and the secondary process with a photoelectron from the photoionization process, X + e X * + e X + e + h e, photoelectron impact excitation...(1.7) In the above four equations, X and Y represents atoms, XY represents a molecule and * represents the excited state. Incoming photons are represented by h and emitted photons are represented as h e. In resonance scattering the exciting and emitting photons are of same frequency. iii) Nightglow: Nightglow occurs during night-times when all direct or Rayleigh scattered sunlight is absent. It occurs when the zenith angle of the sun is greater than ~110. The excitation of nightglow is because of chemical reactions and energy transfer between atoms, molecules and ions. The nightglow emissions are usually very faint with intensities ranging from about 1 to 00 R. The nightglow emanates from two well separated layers in the atmosphere: i) the upper mesosphere and ii) the thermosphere. The prominent nightglow emissions are discussed below with their mechanisms OI nm This redline emission from the atomic oxygen originates at a height of about 50 km in the ionospheric region. The thickness of the emission layer is about 50 km. The intensity of this line is in the range 50 to 100 R. The excitation mechanism for this emission is O + + O O + + O...(1.8) O + + e O + O * ( 1 S 0, 1 D )...(1.9) The excited state designated by the term 1 D is responsible for this emission. The excitation energy is 1.97 ev. 0

21 O * ( 1 D ) O ( 3 P,1) + h (630.0 nm, nm)...(1.10) The red line is a doublet with wavelengths nm and nm. The ratio of the intensities of these wavelengths is 3:1. Hence, only the intensity of the nm line is considered in most of the studies. The life time of the metastable exited state is about 110 s OI nm The first observed emission line in the night airglow was the ( 1 D - 1 S 0 ) forbidden transition in the ground configuration of O at nm, commonly referred to as green line. The intensity of this line is about 50 R. The OI nm emission originates from two distinct layers, one in the F-region around 50 km, the other from a layer of thickness ~10 km centred at ~97 km. In the F-region O ( 1 + S 0 ) is produced by dissociative recombination of O ions as given by the reaction (1.8). The thermospheric contribution in this nightglow accounts for about 0 30%. Fig Simplified energy level diagram of OI green and red lines (adaptedd from Rees, 1989) 1

22 The excited state ( 1 S 0 ) is produced by a two-step process. The emission mechanism is O + O + M O * + M...(1.11) O * + O O + O * ( 1 S 0 )...(1.1) O * ( 1 S 0 ) O * ( 1 D ) + h (557.7 nm)...(1.13) The excitation energy is 4.17 ev and the lifetime of this metastable state is about 0.91 s. The simplified energy level diagram of OI nm and OI nm emissions is shown in Figure OI nm This emission from atomic oxygen comes from the altitudes near the peak height of the F-region. This emission results from direct radiative recombination of O + ions and electrons as shown below. O + + e O * (3p 5 P)...(1.14) O * (3p 5 P) O * (3s 5 S) + h (777.4 nm)...(1.15) The life time of the excited state is s. The intensity of this line is 135 to 00 R. This line results from allowed transition Na Emission The sodium doublet (589.0 nm and nm) is amongst the prominent nightglow emissions. The mesospheric layer of neutral sodium whose origin is believed to be meteor ablation lies between 85 and 105 km, with peak concentration at ~95 km. The sodium airglow emanates from a ~10 km region centred at 90 km. The emission intensity varies

23 significantly with seasons. The excited state is produced through the following oxidationreduction cycle, Na + O 3 NaO + O...(1.16) NaO + O Na * ( P, S) + O...(1.17) followed by the transition, Na * ( P 3/,1/ ) Na( S 1/ ) + h (589.0 nm, nm)...(1.18) Lifetimes of the excited states are approximately 10-8 s and they are allowed transitions OH Meinel Band Emission Vibrationally dis-equilibrated OH radicals make a strong contribution to the airglow in the red end of the visible and IR regions. The emission is due to fundamental ( = 1) and overtone ( > 1) transitions from vibrationally excited radicals in the ground electronic state of OH (X 3/,1/ ). Fundamentals lie in the range from ~800 to ~4500 nm. However, overtones dominate the band and hence the emissions peak in near IR region. The summed intensity of entire Meinel band emissions exceeds 4500 kr. They are intrinsically stronger but relatively difficult to detect because they lie well inside the infra-red region. This emission originates from a layer of about 10 km thickness centred approximately at 87 km. An important reaction that contributes to this emission is H + O 3 OH ( 9) + O + 3.3eV...(1.19) Laboratory investigations show that this reaction is indeed strongly chemiluminescent with OH being the emitter. 3

24 O Band Emission The Herzberg and Chamberlain band emission at far end of the blue and near UV region accounts for about 0.5 kr of nightglow emissions. The most prominent of the O nightglow emissions observed from ground is O atmospheric (0-1) band emission at nm. Its intensity is about 1.5 kr. This emission originates from a layer of ~10 km thickness centered at 94 km. The three body reaction for this emission mechanism is, O + O + M O * + M...(1.0) This emission results from transition between the states, O (b 1 g + ) O ( 3 g - ) + h (atmospheric)...(1.1) Other important reactions are, OH (4) + O O * (b 1 g ) + H...(1.) OH (1) + O O * (a 1 g ) + H...(1.3) O ( 1 D ) + O O * (b 1 g ) + O...(1.4) The lifetime of the metastable b state is about 1 s. The lifetime of the metastable a state is s. The O (0-0) band emission at nm is the most intense emission. Its intensity is ~0 times greater than O (0-1) band emission. This emission cannot be observed from ground due to its total absorption by O molecules below the emission layer. Remote sensing of airglow emissions reveal several important information about the upper mesosphere and ionosphere-thermosphere system which cannot be obtained with other techniques. For example, the remotely sensed emission intensities can be used to derive the concentration of emitting species. The temperatures and neutral winds of the airglow 4

25 emission region could be reliably estimated with high precision interferometry experiments. In addition, from the ratios of the vibrational-rotational line intensities of the molecular band emissions, the kinetic temperature in the upper mesospheric region could be inferred. Further, the waves propagating in the mesospheric region affect the airglow emission intensities by inducing changes in the pressure and temperature of the region [Makhlouf et al., 1998; Swenson and Gardner, 1998; Liu and Swenson, 003]. Similarly, the thermospheric oxygen airglow is sensitive to the fluctuations in plasma density and vertical movements of isoelectron density surfaces. Imaging of the airglow emissions record the column integrated intensity variations over a wide spatial region. Such a technique yields invaluable information about the morphology of gravity waves, nonlinearly evolving waves and instabilities occurring in the mesospheric region and large scale plasma features like plasma depletions and the nighttime equatorward movement of EIA Atmospheric gravity waves The atmosphere is a stably stratified rotating fluid which supports different types of wave motions. Waves occur in the atmosphere whenever the balance between the forces is disturbed by a perturbation and consequently restored. If the perturbation grows without restoration, then instability is said to have formed. In this section, we briefly describe atmospheric gravity waves that are known to play a vital role in the dynamics of upper and middle atmospheric regions. Many excellent texts discuss variety of atmospheric waves in a rigorous manner involving fluid mechanics concepts and assuming conservation of mass (continuity equation), momentum (in both horizontal and vertical directions) and energy [Ex. Beer, 1974; Andrews, 00; Holton, 004]. All the approaches generally assume inviscid 5

26 flow in which the viscosity is negligible, ignoring frictional effects. The role of viscosity or friction will be to damp the waves. Gravity waves are the oscillations of air parcels resulting from the lifting force of buoyancy and the restoring force of gravity. These waves are also known as buoyancy waves. They are transverse in nature and the perturbations associated with them affect the hydrostatic balance. Gravity waves propagate vertically as well as horizontally. The horizontal wavelengths of these waves range from few kilometres to thousands of kilometres and their periods range between few minutes and several hours. Gravity waves occur at all altitudes in the atmosphere and are amongst important dynamical phenomena in all meteorological scales. In the absence of damping, the amplitude of a gravity wave grows with altitude because of exponential decrease in density [Hines, 1960]. As a result, a small perturbation in the lower atmosphere becomes large in the middle and upper atmosphere and these waves play very important role in the energetics and dynamics at those altitudes. In a stably stratified atmosphere, an air parcel displaced vertically upwards will be acted upon by a downward directed gravitational force. In the absence of appreciable friction, the parcel overshoots its equilibrium position and consequently will be acted upon by upward directed buoyant force. In this way, the parcel will undergo an oscillation about its equilibrium position. This characteristic oscillating frequency for a pure vertical displacement of air parcel is called buoyancy frequency or Brunt-Vaisalla frequency and is given by, g dt g g dθ N = T + = dz C...(1.5) P θ dz where, N is the buoyancy frequency (all other symbols were introduced earlier). Square of N is considered to represent static stability of the atmosphere and behave as refractive index for gravity waves. 6

27 It may be noted that the gravity waves are transverse in nature and in which case the buoyancy frequency corresponds to frequency of only those waves that propagate only in the horizontal direction. Other frequencies of gravity waves occur when the air parcel displacements are at an angle to the vertical and are smaller than the buoyancy frequency. Most of the waves propagate at an angle to the horizontal (i.e. displacements are at an angle to the vertical) as shown in Figure The intrinsic frequency of gravity waves () whose air parcel oscillations occur along a plane inclined at an angle to the horizontal is given as, g θ ω = sin β = N sin β...(1.6) θ z From the above equation, it is evident that the largest possible frequency of gravity waves is equal to the buoyancy frequency. The smalles frequency is determined by the coriolis parameter, f=sin, with being the angular velocity of earth and is the latitude. This is because as the frequency of these waves decreases and approaches that of coriolis parameter, their horizontal wavelengths will increase to an extent that the coriolis effect becomes non-negligible. The frequency range of freely propagating gravity waves is thus considered to lie between f and N. The wave parameters such as phase speeds and frequencies observed for the gravity waves are usually known as apparent as the background wind Doppler shifts them. Intrinsic parameters are the true parameters of the waves that are obtained when data on the background winds are available. 7

28 Fig Airparcel displaced from equilibrium position along a plane nclined at an angle to the horizontal (adapted from Nappo, 00) The wave equation for gravity waves is also known as Taylor-Goldstein equation and it can be derived by applying the basic conservation theorems for fluid flow viz. conservation of momentum, mass and thermal energy [For derivation, see Nappo, 00]. In deriving this equation, it is assumed that the longitudinal compressibility and coriolis effects are negligible. These assumptions hold good for substantial portion of gravity wave spectrum. The Taylor-Goldstein equation is given below. d wˆ N + dz ( u c) u 1 1 zz uz + k ˆ = 0 ( ) ( ) 4 w...(1.7) u c H u c H Where, wˆ = w e ˆ1 ( i( kx ωt ) + z / H ) H with ŵ 1 being the perturbation to the vertical wind. In eqn. (1.7), z and x are the Cartesian coordinates representing vertical and horizontal directions, respectively. N is the buoyancy frequency, c is the apparent phase velocity of gravity wave, u is the component of wind along the wave propagation direction, u z and u zz are the first and second derivatives of wind with respect to z direction, H is the scale height and k is the horizontal wave number. 8

29 Note that eqn (1.7) is Helmholtz wave equation for the vertical coordinate, z. The expressions within the bracket in the second term of eqn (1.7) can be equated to the square of the vertical wave number, m, as given below. m N ( u c) u ( u c) 1 u H ( u c) 1 4H zz z = + k...(1.8) Eqn (1.8) gives the relationship between the vertical wave number and intrinsic phase speed, (u-c). It is the dispersion relationship for gravity waves excluding coriolis force and compression effects. The first term of eqn (1.8) is called the buoyancy term and is the dominant term in most of the cases. The second and third terms are called the curvature and shear terms, respectively. The fourth term involving scale height is important for those gravity waves whose vertical wavelengths are of the order of scale height or larger. The fifth term is called the nonhydrostatic term which is merely square of horizontal wavenumber. This term will be of importance only for high frequency gravity waves with small horizontal wavelengths. On assuming that the variations in background wind are not significant, the dispersion relationship can further be simplified as, m N ( u c) 1 4H = k...(1.9) A discussion on relative importance of various terms in dispersion relation is provided in section 3.3, in which we show that for statistical studies concerning gravity wave characteristics eqn (1.9) is sufficient. by eqn (1.30) The phase velocities of gravity waves in horizontal and vertical directions are given 9

30 c = ± ω / k; c = ± ω m...(1.30) x z / However, wave energy transport always occurs at the group velocity, which is given for horizontal and vertical directions as, N( m C gx = u ± ( m + k + + 1/ 4H 1/ 4H ) ) 3/ C Nmk = gz 3/ ( m k +...(1.31) + 1/ 4H ) An important point to note is the opposite directionalities of phase and group velocities in the vertical direction. This indicates that whenever energy is transported upwards by an upward propagating gravity wave, the phase will propagate downwards. Nevertheless, the directions of phase and group propagation are same in the horizontal. Based on their frequencies, gravity waves can be separated into three types as high frequency waves ( ~ N; >> f), medium frequency waves (N >> >> f) and low frequency waves ( << N; ~ f) [Fritts and Alexander, 003]. The high frequency gravity waves usually have short horizontal wavelengths of few kilometres to ~100 kilometers and relatively long vertical wavelengths of ~8 km and above. Their time periods range from few minutes to about an hour or two. These waves are believed to carry large flux of energy and momentum upwards. When the vertical wavelength becomes large and m 0, the wave will undergo total internal reflection where the vertical group velocity can change sign. Such a level is called turning level or level of reflection. In section 3.3, we will show that such a situation often arises for high frequency waves when they propagate opposite to the mean background wind. Above this level, the wave becomes evanescent. Evanescent waves are standing modes without vertical propagation and their 30

31 frequencies will be less than the local buoyancy frequency. If turning levels occur both above and below, the wave is said to be ducted in that region. Short horizontal wavelengths are easily reflected and trapped in ducts. Once ducted, the wave could travel large horizontal distances of few thousands of kilometres, depending upon the length of the duct channel. This enables the ducted waves to effectively carry energy from source region to distant regions where they undergo breaking. The dispersion relation for high frequency wave cannot be further simplified than that given in eqn. (1.9). However, it can be written in the other form as, m k N 1 4H = k ω...(1.3) We prefer the form given in eqn. (1.9) as it is simpler to use in calculations involving phase speeds instead of intrinsic wave frequency. Waves with time periods of few hours to ~8 hours are generally referred to as medium frequency waves. Their horizontal wavelengths lie in the range of few hundred kilometres. Most of the gravity waves observed by radars are medium frequency gravity waves. For these waves, the dispersion relationship can be further simplified as, N k N =...(1.33) m = ( u c) ω This relation provides valuable insights of gravity wave properties and the effects induced by changes in the background wind and stability. It may be noted that the vertical wavelength is proportional to the intrinsic phase speed. Therefore, as the background wind approaches the apparent phase speed of the wave, the vertical wavelength approaches zero. Such a level is called critical level where the wave will be absorbed by the mean background flow and the vertical propagation past the level is significantly affected. Instabilities and 31

32 dissipation mechanisms act to damp the waves as the latter approach the critical level, thereby extracting energy from the waves before they reach critical level. Further, the vertical group velocity decreases faster than the phase velocity as the wave approaches the critical level implying that the energy propagation is hindered more effectively. Gravity waves with large horizontal wavelengths of thousands of kilometres and time periods of several hours (typically >8 hours) are called low frequency waves. Usually their vertical wavelengths will be less than few kilometres. Since their horizontal wavelengths are large, they are affected by the rotation of the earth. The dispersion relationship for the low frequency gravity waves is thus written as, m k N =...(1.34) ω f Another insight that could be obtained from this dispersion relation is that the vertical wavelength becomes zero (m) when equals f. Therefore, any wave approaching a critical level is Doppler shifted to lower intrinsic frequencies until it reaches f. The level where the wave intrinsic frequency,, reaches f will occur somewhere below the level at which u = c. This implies that the true critical level is actually the level at which = f. The low frequency gravity waves are also known as inertia gravity waves. It is important to gather information about the sources of gravity waves and their global distribution. Further, it would be interesting to know whether the characteristics of waves vary depending on their sources. A variety of source mechanisms are proposed for the generation of gravity waves. It appears all these mechanisms contribute to the spectrum of the observed waves. Here, we briefly discuss important source mechanisms. Topographic generation of gravity waves is among the vastly studied source mechanism [Nappo, 00; Fritts and Alexander, 003]. The flow of wind across terrain 3

33 elevations or depressions such as mountains, ridges, hills, canyons and valleys can give rise to generation of gravity waves. The gravity waves generated are also known as mountain waves or lee waves. The mountain waves are standing wave patterns with zero intrinsic phase speeds generated by uplift of air across a mountain. The lee waves are ducted waves formed in the leeward side of mountains. They propagate only in horizontal direction and are trapped close to the ground. Hence, lee waves do not contribute to wave processes occurring in the upper atmosphere. Convective generation is believed to be the most important source for gravity waves observed in middle and upper atmosphere, especially over the tropics. Convective activity could give rise to waves with high intrinsic phase speeds and it is also capable of generating a large spectrum of waves. However, the exact mechanisms by which gravity waves are generated through convection remains unclear. Three major mechanisms proposed are: (i) pure thermal forcing associated with the latent heat release of convective complexes, (ii) obstacle effect in which the rising convective element acts as an obstruction to the horizontal flow thereby generating gravity waves and (iii) mechanical oscillator effect in which intense updrafts and downdrafts within the convective complexes are believed to generate gravity waves. Another possibility closely associated with the mechanical osciallator effect is the periodic overshooting of the tropopause by means of intense convection that could trigger waves at the lower stratospheric altitudes. Shear generation of gravity waves has been studied for many years but still remains to be a least understood source. Intense shears resulting in unstable flow are proposed to generate gravity waves. Shear instabilities will be generated in the regions of intense shear. Those instability modes appear to involve in non-linear interaction between the propagating modes thereby rapidly exciting the propagating modes. These propagating modes emanate as gravity waves from the regions of intense shear and the process is also referred to as 33

34 envelope radiation [see Fritts and Alexander, 003 and references cited therein]. Shear generation is capable of exciting high frequency gravity waves too. The gravity waves produced in the vicinity of jet streams are believed to fall in this category. Other proposed source mechanisms are wave-wave interactions, generation of secondary waves from breaking of primary waves, geostrophic adjustment and eclipse cooling. In addition, at high latitudes auroral heating may play a role. Of these, wave-wave interactions and generation of secondary waves from breaking of primary waves have the potential to generate waves of high and medium frequencies. Gravity waves are observed with a variety of observational techniques like optical measurements, lidar measurements, radar observations, radiosonde profiles, rocket soundings, aircraft measurements, microbarographs, sodars, limb viewing measurements made by satellites, etc. Most of the abovementioned observations yield information about the apparent periodicities of the waves from time series data or vertical structure of the wave from altitude profiles of measured parameters. Further, they are sensitive to the medium and low frequency spectrum of gravity waves. In order to study high frequency waves, observations of the parameters like wind, temperature, airglow intensity, etc., should be made in high spatio-temporal resolution. Among these measurement techniques, a unique method capable of providing direct information on the horizontal structure of the high frequency gravity waves in the upper mesospheric region is the all-sky airglow imaging technique. The airglow images capture the two dimensional horizontal wave field affecting the airglow layers directly. By analyzing images of more than one upper mesospheric airglow emissions, it is possible to infer the three dimensional structure of the waves. Recently, attempts were being made to obtain the three dimensional information of the wave field by means of observing the same region of the sky with multiple cameras, a method similar to triangulation [Nygren et al., 1998; Moreels et al., 008]. 34