Observations of Water Vapour in the Middle Atmosphere

Observations of water vapour in the middle atmosphere

Stefan Lossow

AKADEMISK AVHANDLING för filosofie doktorsexamen vid Stockholms Universitet att framläggas för offentlig granskning den 19:e september 2008 Observations of water vapour in the middle atmosphere Doctoral thesis Stefan Lossow Cover illustration arranged by Stefan Lossow. Odin picture by Bill Frymire/Pressens Bild and Swedish Space Corporation. Courtesy of Swedish Space Corporation.

ISBN 978–91–7155–721–6, pp. 1 – 63 c Stefan Lossow, Stockholm 2008

Stockholm University Department of Meteorology 10691 Stockholm Sweden

Printed by Universitetsservice US-AB Stockholm 2008 In memory of my grandmother

† 2006

Contents

Abstract 7

List of papers 9

1 Introduction 11 1.1 Structure of the atmosphere ...... 11 1.2 Middle atmosphere dynamics ...... 13 1.3 Water vapour in the middle atmosphere ...... 18 1.3.1 General distribution ...... 20 1.3.2 Recent research ...... 23

2 Odin 25 2.1 Satellite-borne water vapour measurements in the middle atmosphere ...... 25 2.2 The Odin satellite ...... 28 2.3 SMR measurements of the 557 GHz band ...... 29 2.4 Proﬁle retrieval ...... 31

3 Hygrosonde 33 3.1 Measurement principle ...... 34 3.2 Instrument aspects ...... 35

4 Results of this thesis 37

5 Outlook 41

5 Contents

Acknowledgements 45

References 48

6 Abstract

Water vapour is the most important greenhouse gas and plays a fundamental role in the climate system and for the chemistry of the Earth’s atmosphere. This thesis presents observations of water vapour in the middle atmosphere with a particular focus on the mesosphere. The majority of these observations presented in this thesis have been performed by the Swedish satellite Odin, providing global observations since 2001. Further observations come from the Hygrosonde-2 campaign in December 2001 based on balloon and rocket-borne measurements. A general overview of Odin’s water vapour measurements in the middle atmosphere is given. The optimisation of the mesospheric water vapour retrieval is presented in detail. The analysis of the observations has focused mainly on different dynamical aspects utilising the characteristic of water vapour as a dynamical tracer in the middle atmosphere. One application is the mesospheric part of the semi- annual oscillation (SAO). The observations reveal that this oscillation is the dominant pattern of variability between 30◦S and 10◦N in the mesosphere up to an altitude of 80 km. Above 90 km the SAO is dominating at all latitudes in the tropics and subtropics. It is shown that the SAO exhibits a distinct phase change between 75 km and 80 km in the tropical region. This thesis also presents the first satellite observations of water vapour in the altitude range between 90 km and 110 km, extending the observational database up into the lower thermosphere. In the polar regions water vapour exhibits the annual maximum during winter time above 95 km, mainly caused by upwelling during this season. This behaviour is different from that observed in the subjacent part of the mesosphere where the annual maximum occurs during summer time. The Hygrosonde-2 campaign provided a high resolution measurement of water vapour in the vicinity of the polar vortex edge. This edge prevents horizontal transport causing different water vapour characteristics inside and outside the polar vortex. The observations show that this separating behaviour extends high up into the mesosphere. Small scale transitions in the Hygrosonde-2 profile between conditions inside and outside the vortex coincided with wind shears caused by gravity waves.

List of papers

This thesis consists of an introduction and the following papers:

(1) Urban, J., N. Lautié, D. Murtagh, P. Eriksson, Y. Kasai, S. Lossow, E. Dupuy, J. de La Noë, U. Frisk, M. Olberg, E. Le Flochmoën and P. Ricaud; Global observations of middle atmospheric water vapour by the Odin satellite: An overview; Planetary and Space Science; 55; 1093 – 1102; 2007 (2) Lossow, S., J. Urban, P. Eriksson, D. Murtagh and J. Gumbel; Critical parameters for the retrieval of mesospheric water vapour and temperature from Odin/SMR limb measurements at 557 GHz; Advances of Space Research; 40; 835 – 845; 2007

(3) Lossow, S., J. Urban, J. Gumbel, P. Eriksson and D. Murtagh; Observa- tions of the mesospheric semi-annual oscillation (MSAO) in water vapour by Odin/SMR; Atmospheric Chemistry & Physics Discussion; 8; 10153 – 10187; 2008; revised (4) Lossow, S., J. Urban, H. Schmidt, J. Gumbel, P. Eriksson and D. Murtagh; Winter time water vapour in the polar upper mesosphere and lower thermosphere - First satellite observations by Odin/SMR; Manuscript (5) Lossow, S., M. Khaplanov, J. Gumbel, J. Stegman, G. Witt, P. Dalin, S. Kirkwood, F. J. Schmidlin, K. H. Fricke and U. A. Blum; Middle atmospheric water vapour and dynamics in the vicinity of the polar vortex during the Hygrosonde-2 campaign; Atmospheric Chemistry & Physics Discussion; 8; 12227 – 12252; 2008

Chapter 1

Introduction

1.1 Structure of the atmosphere

The atmosphere can be divided into different altitude regions based on various criteria. A common way is to use the temperature structure, which is shown in figure 1.1 In this case the transition layers (”pauses”) between these different altitude regions are defined by a reversal of the temperature gradient. The lowermost layer, the troposphere (greek: tropos = turning, mixing), ranges up to about 8 km in the polar regions and up to about 17 km at the equator. In this layer the everyday weather is taking place. The troposphere is characterised by a temperature decrease of about 6.5 K/km. At the top of the troposphere, i.e. at the tropopause, temperatures between 200 K (at the equator) and 230 K (at the poles) can be observed. The tropospheric energy input arises mainly from thermal infrared radiation emitted from the Earth’s surface and latent heat release. The layer on top of the troposphere is the stratosphere (greek: strato = layered). It extends up to an altitude of about 50 km. In the lowest part of the stratosphere the temperature is often nearly isotherm. Above about 20 km the temperature increases rapidly and reaches a maximum approximately at an altitude of 50 km. This increase is caused by the absorption of ultraviolet solar radiation (mainly in the wavelength range 200 nm to 310 nm) by the ozone layer. 90 % of the total atmospheric ozone can be found in the middle stratosphere [Brasseur and Solomon, 1998]. Due to the temperature increase with altitude the stratosphere is, unlike the troposphere, stably stratified. In the mesosphere (greek: mesos = middle) the temperature decreases once again with altitude, typically with a gradient of about (3±1) K/km. The mesopause is the boundary between the mesosphere and the thermosphere

11 Chapter 1: Introduction

110 Thermosphere 100 NLC 90 Rocket 80

70 Mesosphere

50 Altitude [km] 40

30 Stratosphere Radiosonde

20 PSC

10 Airplane Troposphere Mesopause 0 100 120 140 160 180 200 220 240 260 280 300 320 Temperature [K] Figure 1.1: The temperature structure of the Earth’s atmosphere. The black profile represents conditions at the equator while the grey profile describes the conditions during summer in the polar region. and is defined as the small layer where the local temperature minimum occurs above 80 km. The mesopause tends to occur at two preferred altitude levels, i.e. at around 100 km (normal state) and at around 88 km (polar summer state) [von Zahn et al., 1996; Lübken, 1999]. The mesopause is the coldest place in the Earth’s atmosphere, where temperatures can drop below 130 K during polar summer [Lübken, 1999]. In the thermosphere the temperature increases rapidly and can reach values of more than 1500 K. This increase is caused by the absorption of high energy solar radiation in the wavelength region below 200 nm. One would not feel these extremely high temperatures due to the ”near vacuum” conditions in this altitude region. The mean free path of the remaining atoms and molecules is in the order of metres, hence heat exchange events occur only rarely. The altitude range between 10 km and 100 km is usually referred to as the middle atmosphere, essentially comprising the stratosphere and mesosphere. This part of the atmosphere is an interesting region for atmospheric science and understanding is still rather limited especially in the upper half, on which this thesis is mainly focused. The middle atmosphere can be clearly distinguished from the troposphere below and thermosphere above in many

12 Chapter 1: Introduction aspects, e.g. in characteristics and processes that matter for radiation, chemistry and dynamics. This is also valid for water vapour, the main topic of this thesis.

1.2 Middle atmosphere dynamics

Dynamics in the atmosphere is mainly driven by the heat differences induced due to the inhomogeneous insolation. In the middle atmosphere an additional important mechanical driving component arises from the angular momentum deposition of breaking waves. The troposphere is thermically adjusted by convection in the tropics, while in the extra-tropics the adjustment is based on baroclinic instabilities. In the middle atmosphere the thermal inertia is determined by the radiative relaxation time. This relaxation time is typically not longer than a few weeks, so that the radiative equilibrium temperatures in the middle atmosphere follow the Sun. The upper panel of figure 1.2 shows the zonal mean temperature distribution for July as a function of altitude. From radiative balance considerations right after the northern hemisphere summer solstice, maximum temperatures can be expected in the northern hemisphere polar region while minimum temperatures are expected to be observed in the southern hemisphere polar region. Obviously this picture works well in the stratosphere and lower mesosphere. However, higher up in the mesosphere the situation is vice versa with lower temperatures over the summer pole than over the winter pole. This is a consequence of a mechanically driven circulation, as it will be discussed later. In a rotating fluid like the Earth’s atmosphere the meridional temperature gradients can be sustained, since the arising pressure gradient is balanced by the Coriolis force. These meridional temperature gradients have a direct consequence for the zonal wind distribution, as expressed by the thermal wind relation: ∂u R ∂T f · = − air · (1.1) ∂z rEarth · H ∂φ where f is the Coriolis parameter, u the zonal wind, z the altitude, Rair the gas constant for air, rearth the radius of the Earth, H the scale height and φ the latitude. According to this relation the vertical gradient of the zonal wind is proportional to the meridional temperature gradient. The lower panel of figure 1.2 shows the zonal mean distribution of the zonal wind, again for July conditions. From the thermal wind relation the zonal wind distribution in the middle atmosphere can quickly be derived. In the northern (summer) hemisphere a positive meridional temperature gradient can be observed in the stratosphere and lower mesosphere. At the same time

13 Chapter 1: Introduction

July zonal mean temperature [K] 100 180 200 180 90 140 160 80 200 220 70 200 240 220 60 240 260 260 50 260 40 260

Altitude [km] 240 240 30 220 200 20 200 220 10 220 240 220 240 280 260 0 260 −75 −60 −45 −30 −15 0 15 30 45 60 75 Latitude [degree]

July zonal mean zonal wind [m/s] 100

0 40 90 20 60 0 0 80 0 −20 40 70 60 40− 20 − 60 20 − 40 20 60 − 50 0

− 20 40 20 Altitude [km]

30 60 20 40 0 20 20 10 40 0 0 0 0 0 −75 −60 −45 −30 −15 0 15 30 45 60 75 Latitude [degree] Figure 1.2: The zonal mean distribution of temperature (upper panel) and zonal wind (lower panel) for July from HAMMONIA (Hamburg Model of the Neutral and Ionized Atmosphere, Schmidt et al. [2006]). Negative zonal wind values describe easterlies (westward winds), while positive values describe westerlies (eastward winds). Courtesy of Hauke Schmidt, MPI for Meteorology, Hamburg, Germany. 14 Chapter 1: Introduction the Coriolis parameter is positive in the northern hemisphere, resulting in a decrease of the zonal wind wind with altitude. Consequently the zonal wind decreases in the northern hemisphere above the tropopause and easterlies develop. In the southern (winter) hemisphere the meridional temperature gradient is also positive. However, in this hemisphere the Coriolis parameter is negative, resulting in an increase of the zonal wind with altitude. In accordance, strong westerlies can be observed in the winter hemisphere. Higher up in the mesosphere, the meridional temperature gradients reverse and in correspondence also the vertical gradients of the zonal wind. This leads to a deceleration of the easterlies and westerlies and can even result in a reversal of the zonal wind direction as observed in the summer mesospause region. The zonal flows are usually strongest in the atmosphere. Meridional flows on the other hand are in general rather weak and require some kind of forcing. This is where the mechanical forcing by breaking waves comes in. When these waves break they deposit their angular momentum and insert a wave drag. This wave drag forces meridional flows and causes the middle atmosphere to deviate from the radiative equilibrium, as in the mesosphere shown in the upper panel of figure 1.2. Two types of waves are especially important for the mechanical forcings in the middle atmosphere, namely planetary-scale Rossby waves and smaller scale gravity waves. Both types are in general excited in the troposphere and carry their momentum upwards. Rossby waves deposit their momentum in the stratosphere, while gravity can propagate through into the mesosphere, where the usually break and deposit their momentum. Rossby waves are excited in the troposphere by land-sea contrasts and have the Coriolis force as restoring force. These waves can only propagate in westerly conditions, limiting their propagation to the winter hemisphere. When the Rossby waves dissipate in the stratosphere they exert always a negative wave drag. This drag causes a deceleration of the westerlies which forces a poleward flow. Due to reasons of mass conservation this poleward flow is accompanied by upwelling in the tropics and downwelling in the extra-tropics. This circulation cell is referred to as the Brewer-Dobson circulation, first observed in the distribution of long-lived trace gases [Brewer, 1949]. Since the Rossby waves only can propagate in the winter hemisphere the Brewer-Dobson circulation is mostly a winter time phenomenon. The Brewer-Dobson circulation is stronger in the northern hemisphere as the land-sea contrasts generate stronger Rossby waves in this hemisphere. This corresponds to stronger down-welling and higher temperatures due to adia- batic heating in the Arctic winter time stratosphere as compared to Antarc- tic conditions, which has significant consequences for the ozone destruction

15 Chapter 1: Introduction in the polar vortex. Gravity waves on the other hand are excited by several mechanisms, as the air flow over mountains, frontal systems, convection or non-linear wave- wave interactions. The restoring force of these waves are buoyancy or gravity. Unlike Rossby waves gravity waves can exert both a positive or negative wave drag upon the atmosphere where they break. Gravity waves encounter different propagation conditions during summer and winter based on their filtering by the background winds. Filtering occurs when the zonal phase speed of the waves equals the zonal wind speed of the background wind. In the winter time, when westerly conditions prevail, the majority of gravity waves with positive phase speeds are filtered out, so that mainly gravity waves with negative phase speeds are allowed to propagate up into the mesosphere. When these waves break they cause a negative wave drag, which leads to a weakening of the westerlies and forces a poleward flow. During summer time the conditions are vice versa. The easterlies in the middle atmosphere allow mainly gravity waves with positive phase speeds to propagate up into the mesosphere. Now, when these waves break they induce a positive wave drag, which causes a weakening of the easterlies and drives a equatorward flow. Combined, this results in a solstitial pole-to-pole flow in the mesosphere from the summer to the winter hemisphere. Again, due to the requirement of mass conservation this pole-to-pole flow is accompanied by upwelling over the mesospheric summer pole and downwelling over the winter pole. This circulation drags the atmosphere so strongly away from the radiative equilibrium that the temperatures in the summer mesopause region are actually lower than in polar night conditions, as it can be seen in figure 1.2 [Holton, 1982; Shepherd, 2000]. The following list includes a short overview of some more important dynamical features in the middle atmosphere related to this thesis.

• Quasi-biennal oscillation (QBO) The QBO describes the zonal wind oscillation of easterly and westerly regimes descending from the tropical middle stratosphere towards the tropopause, which was ﬁrst observed in the 1950s [Graystone, 1959; Ebdon, 1960; Reed et al., 1961]. Having a mean period of about 28 months (can vary between 20 and 36 months) the QBO is not connected to any geophysical periodicity. The phenomenon is approximately symmetric around the equator with a meridional half width of about 12 degrees latitude. Maximum amplitudes can be observed in the altitude range between 30 hPa (∼23.5 km) and 10 hPa (∼31 km), typically twice as big during the easterly phase as compared to the westerly phase. Initially it was thought that equatorially trapped Kelvin waves provide the westerly momentum to drive the QBO,

16 Chapter 1: Introduction

while Rossby gravity waves provide the easterly momentum [Lindzen and Holton, 1968; Holton and Lindzen, 1972]. However observations showed that the forcing by these waves is not suﬃcient to drive the QBO with the observed amplitude and period. The additional forcing needed is provided by momentum deposition of a broad spectrum of gravity waves [Baldwin et al., 2001, and references therein]. The QBO is accompanied by anomalies in temperature, trace gases and other dynamical variables. The eﬀects of the QBO are not limited to the tropical stratosphere, they play a major role for the variability also in the extra-tropics [e.g. Holton and Tan, 1980; Dunkerton et al., 1988; Holton and Austin, 1991; Labitzke and van Loon, 1999; Wu et al., 2008a].

• Semi-annual oscillation (SAO) Higher up, above 35 km, the semi-annual oscillation (SAO) is the dominant pattern of variability in the tropical region. This oscillation describes a semi-annual variation, that has first been observed in zonal wind and temperature data [Reed and Rogers, 1962; Reed, 1966]. The SAO peaks at the stratopause and the upper mesosphere with a minimum in-between occuring at about 65 km. In accordance a dis- tinction between stratospheric (SSAO) and mesospheric semi-annual oscillation (MSAO) is made. The SSAO and MSAO variations are approximately 180◦ out of phase with respect to each other [Hirota, 1980; Garcia et al., 1997]. The forcing mechanisms of these two types of SAO are rather different. The easterly phase of the SSAO in zonal wind is due to the meridional advection of easterlies from the summer hemisphere across the equator (see lower panel of figure 1.2) and the momentum deposition of breaking planetary waves. The westerly phase is forced through the momentum deposition of ultra-fast Kelvin and internal gravity waves [e.g. Holton and Wehrbein, 1980; Hitchman and Leovy, 1988; Ray et al., 1998]. The forcing of the MSAO is thought to be due to a broad spectrum of gravity waves and high-speed Kelvin waves excited in lower atmosphere propagating vertically upwards. These waves are filtered by the zonal wind variations of the SSAO, allowing largely only waves with the opposite zonal propagation direction as the zonal wind direction to propagate higher up into the middle and upper mesosphere, where they usually break. The wave filtering by the SSAO in the zonal wind accounts for the observed phase shift between the SSAO and MSAO variations [Dunkerton, 1982].

• Atmospheric tidal waves In the upper part of the middle atmosphere atmospheric tidal waves

17 Chapter 1: Introduction

exhibit substantial amplitudes so that these waves play an important role in governing the dynamics of this region. Atmospheric tides are thermally driven by the periodic absorption of solar radiation, mainly the absorption of infrared radiation by water and water vapour in the troposphere and the absorption of ultraviolet radiation by stratospheric ozone. Another important driver is the latent heat stored in water vapour and transported by the weather systems in the troposphere. They have periods that are harmonics of 24 h and typical vertical wavelengths of 25 km and more. There are two types of atmospheric tides, i.e. migrating and non-migrating tides. Migrating tides move westward with the apparent motion of the Sun. Non-migrating tides can propagate both westward and eastward or they are station- ary. Detailed information about tides can be found for example in Chapman and Lindzen [1970]; Burrage et al. [1995b, a]; Hagan et al. [1997a, b]; Hagan and Forbes [2002, 2003]; Murphy et al. [2006]; Wu et al. [2008a, b].

1.3 Water vapour in the middle atmosphere

Water vapour is a key constituent in the Earth’s atmosphere. First and foremost it is the most important greenhouse gas. At the Earth’s surface it contributes to bearable global mean temperatures of about 288 K (15◦C) as compared to 255 K (-18◦C) without any greenhouse effect. Thus any change in atmospheric water vapour will have important implications for the global climate. The effect of such a change strongly depends on the altitude where this change occurs. The most significant impacts are caused by humidity changes in the upper troposphere/lower stratosphere region (UT/LS). Esti- mations have shown that a 12 % to 25 % increase of the upper tropospheric humidity results in effects that correspond to a doubling of carbon dioxide (CO2). In addition water vapour is also part of several potentially important global warming feedback mechanisms. Increasing temperatures at the Earth’s surface will cause an increase in water vapour due to enhanced evaporation. The direct radiative effect of such water vapour increase will furthermore increase the surface temperatures. This represents a positive global warming feedback mechanism. Still the net sign of this feedback is under debate, since also effects of accompanied cloudiness changes have to be considered. In the middle atmosphere long-wave emissions of water vapour in the mid- and far-infrared contribute a small part to the cooling of the atmosphere besides the cooling from CO2 and O3.

Water vapour is the primary source for hydrogen radicals HOx (OH, H, HO2) in the middle atmosphere. These radicals are involved in the catalytic

18 Chapter 1: Introduction cycles that destroy ozone with their contribution dominating above 50 km. Very low temperatures inside the polar vortex allow the formation of polar stratospheric clouds (PSC), especially in Antarctic. These clouds usually occur in the altitude range between 15 km and 25 km. On the surface of their ice crystals heterogeneous chemical processes lead to a release of active chlorine species, which are responsible for the severe ozone destruction. The extremely low temperatures observable in the summer mesopause region allow marginally the heterogeneous formation and growth of ice particles, despite the low water vapour concentrations in this region. The ice particles are formed near the mesopause around 88 km. The particles continue to grow by direct deposition of the ambient water vapour onto their surface and sediment. These ice particles give rise to fascinating phenom- ena like noctilucent clouds (NLC1) and polar summer mesosphere echoes (PMSE) [e.g. Thomas, 1991; Rapp and Thomas, 2006; Rapp and Lübken, 2004]. When the ice particle radii exceed ∼30 nm they scatter light effi- ciently such that they might be observed optically as NLC with the naked eye, by ground-based or space-borne optical instruments. NLC occur typically in the altitude range between 82 km to 86 km. Below, the ice particles encounter warmer temperatures and evaporate rapidly. PMSE are radar echoes, observable over a broad frequency range, which are mainly based on the smaller, subvisible ice particles. All the ice particles are immersed in the plasma of the D-region. Hence, electrons attach to the ice surfaces so that the particles become charged. In addition, the 80 km to 90 km altitude range is the region in the atmosphere where gravity waves propagating from below grow unstable and produce turbulence. The charged ice particles are transported by the turbulent velocity field leading to small scale structures in the spatial distribution of the charged particles and, because of charge neutrality requirements, to small scale structures in the spatial distribution of the electron number density. Hence, the transport of charged ice particles by the turbulent velocity field leads to the occurrence and maintenance of irregularities in the radio refractive index (which is effectively determined by the electron number density at these altitudes) which can be observed by radars on the ground as PMSE. In the middle atmosphere water vapour is also a particularly convenient tracer for dynamics, since its chemical lifetime is comparable to typical time scales of dynamical processes. The chemical lifetime of water vapour is several years in the lower stratosphere, months around the stratopause and of the order of a few days at 100 km [Brasseur and Solomon, 1998]. The tracer characteristic of water vapour is utilised in a major part of this thesis (see paper (3), (4) and (5)).

1also called polar mesospheric clouds (PMC)

19 Chapter 1: Introduction

1.3.1 General distribution

At the Earth’s surface water vapour constitutes in average 1.3 % of the atmospheric composition, in the tropical region it can be more than 3 %. In the troposphere the amount of water vapour decreases rapidly with increasing altitude. In the middle atmosphere only a very small fraction of the surface water vapour volume mixing ratios can be observed. Water vapour enters the stratosphere primarily by ascent through the tropical tropopause transition layer (TTL), secondarily by poleward transport along isentropic surfaces which span both the uppermost troposphere and lowermost stratosphere [Holton et al., 1995]. The TTL acts as a so called ”cold trap” for water vapour. Due to the negative temperature lapse rate in troposphere and a high altitude the tropical tropopause exhibits very low temperatures (see upper panel of figures 1.2). These cold temperatures inside the TTL cause a strong reduction of water vapour due to freeze-drying. The ice particles formed during this process sediment subsequently. The exact mechanisms behind this dehydration, the locations where the air parcels attain the final dehydration upon their entry into the stratosphere and where they finally enter the stratosphere is a major topic of current research. In general the Western Pacific area plays an important role in this context, during boreal summer also the deep Monsoon over Tibet is a key pathway of air into the stratosphere. The water vapour throughput into the stratosphere has a seasonal variation according to the TTL temperature. Coldest and dri- est tropopause conditions occur during boreal winter (December to March), warmest and moistest conditions can be observed between June and Septem- ber. This seasonal variation of the water vapour throughput at the tropical tropopause is referred to as the tape recorder effect [Mote et al., 1996]. The mean throughput is around 3.7 ppmv [e.g. Kley et al., 2000]. The TTL is part of the Brewer-Dobson circulation, which transports the long-lived water vapour upon its entry into the stratosphere towards stratospheric mid-latitudes and further to the polar area where it descends. The blue profile in figure 1.3 shows the general altitude distribution of water vapour in the middle atmosphere. The water vapour concentration increases in the stratosphere, reaches a maximum around the stratopause and decreases continuously in the mesosphere. In the stratosphere additional water vapour is formed by the irreversible oxidation of methane (reaction 1.2).

CH4 + 2 · O2 → 2 · H2O + CO2 (1.2)

One methane molecule is in principle converted into two water vapour molecules. In reality the eﬃciency is a little bit smaller because the radicals occurring immediately can also form molecular hydrogen (H2). This con-

20 Chapter 1: Introduction

100

50 Altitude [km]

10 0 1 2 3 4 5 6 7 8 Water vapour [ppmv] Figure 1.3: Examples of typical water vapour proﬁles in the middle atmosphere: general distribution (blue), at the equator around equinox (red) and in the polar region in summer time (green). The proﬁles have been adapted from measurements of the HALOE (Halogen Occultation Exper- iment) instrument aboard UARS (Upper Atmosphere Research Satellite) and extended in the altitude range 85 km to 100 km with model results. version is completed somewhere in the altitude range between 60 km and 70 km. The main sink process of water vapour in the stratosphere is the reaction with O(1D) (reaction 1.3). The water vapour production outweighs the sink process, consequently the water vapour concentration increases in the stratosphere. With increasing altitude the destruction of water vapour by photodissociation (reaction 1.4) becomes more and more important. In the lower mesosphere this is caused by radiation in the Schumann-Runge bands (λ=176 nm - 193 nm).

1 H2O + O( D) → 2 · OH (1.3)

H2O + h · ν → H + OH (λ < 200 nm) (1.4)

A balance between these source and sink processes usually occurs around the stratopause, resulting in a water vapour maximum in this altitude range. This water water maximum is sometimes referred to as the ”conventional” water vapour maximum. Inside the polar vortex however this maximum

21 Chapter 1: Introduction may be observed as much as 20 km lower in altitude, due to the prevailing subsidence in the vortex area. This is addressed in the last paper of this thesis. In the mesosphere, water vapour decreases in general with altitude, since additional sources are usually missing. Above about 70 km water vapour is primarily destroyed due to photodissociation by solar Lyman-α radiation (λ=121.6 nm). In the mesosphere two exceptions from the general water vapour distribution described above can be found, limited in time and space. In polar areas in summer and in the tropics around equinox an additional water vapour maximum might be observed in the altitude range between 65 km and 75 km [Nedoluha et al., 1996; Summers et al., 1997; Seele and Hartogh, 1999]. This can be seen in the red and the green proﬁle shown in ﬁgure 1.3, represent- ing such equatorial and polar area conditions respectively. This maximum is formed by an interplay of chemistry and transport during strongest insolation [Sonnemann et al., 2005]. During such conditions the production of hydrogen radicals is strongest, which are essential for an autocatalytic production of water vapour from the hydrogen reservoir. The two most important net reactions of this production are:

H2 + O → H2O (1.5)

H2 + O3 → H2O + O2 (1.6)

This production occurs slowly below a cross-over height of about 65 km. Above this height, water vapour photodissociation dominates and hydrogen is formed (see reaction 1.4). Below the cross-over height the vertical winds need to be weak to allow a sufficient time for the water vapour production. Above the cross-over height however the vertical winds need to be strong to transport the water vapour enriched air upwards and to form the maximum, before the water vapour destruction by photodissociation takes over again. Another maximum might be observed in a small layer around 82 km in the polar area during summer time, as indicated in the green profile in figure 1.3. This peak is caused by the redistribution of water vapour by the ice particles forming PMSE and NLC. These ice particles eventually sediment to altitudes where they encounter warmer temperatures (≥ 150 K) and subsequently sublimate rather rapidly in the described small layer [Summers et al., 2001; von Zahn and Berger, 2003]. Since water vapour is a dynamical tracer in the middle atmosphere its seasonal variation is mainly connected to the large scale dynamics. In the polar region the large-scale upwelling in summer and subsidence in winter time result in annual cycle of the water vapour concentration. An annual cycle can also be observed in the the lowermost polar stratosphere where

22 Chapter 1: Introduction the seasonal variation is controlled by transport of water vapour from the TTL towards higher latitudes via the Brewer-Dobson circulation and along isentropic surfaces [Pan et al., 2000; Nedoluha et al., 2002]. Towards the equator the dominant annual variation gradually changes from an annual cycle towards a semi-annual cycle (see also paper (3)) [Randel et al., 1998; Jackson et al., 1998]. An annual cycle only dominates in the equatorial area below about 35 km, where the tape recorder signal in water vapour is transported upwards from the TTL.

1.3.2 Recent research

Recent research on the middle atmospheric water vapour has been and is dominantly focusing on the understanding of the troposphere-stratosphere exchange processes of water vapour [e.g. Kley et al., 2000; Sherwood and Dessler, 2000; Fueglistaler et al., 2004; Bonazzola and Haynes, 2004; Engel et al., 2006; Lelieveld et al., 2007; Gettelman et al., 2008]. At the turn of the millennium several studies reported about an increase of the stratospheric water vapour concentration [Oltmans et al., 2000; Michelsen et al., 2000; Rosenlof et al., 2001]. More recent results indicate that this increase has ceased and that over a period of about 3 to 4 years actually a decrease could be observed [Nedoluha et al., 2003; Randel et al., 2004]. Ozone depletion inside the polar vortex remains an important topic and correspondingly also the influence of water vapour on this destruction by the formation of polar stratospheric clouds [e.g. Rex et al., 2004; Eyring et al., 2007]. As a consequence continuous monitoring and improved understanding of changes in stratospheric water vapour and water vapour entering the stratosphere are of the greatest importance. The focus of mesospheric water vapour research is mainly on the budget in the polar summer mesopause region. This comprises the concentration of water vapour in the presence of NLC [Summers et al., 2001; von Zahn and Berger, 2003; Russell et al., 2008], inter-hemispheric differences [Hervig and Siskind, 2006; Lübken and Berger, 2007] and possible short-term trends [von Zahn et al., 2004]. Furthermore the size of the solar influence on the mesospheric water vapour distribution [Chandra et al., 1997; Sonnemann and Grygalashvyly, 2005], 5-day planetary wave signatures in water vapour [Sonnemann et al., 2008] and possible global long-term trends due to the increase of methane [Grygalashvyly et al., 2007] have been addressed.

23 24 Chapter 2

Odin

2.1 Satellite-borne water vapour measurements in the middle atmosphere

Satellite-borne measurements have played and continue to play a major role in improving our knowledge and deeper understanding of the water vapour distribution and its variability in the middle atmosphere. Various techniques and different spectral ranges have been used to retrieve water vapour in the middle atmosphere. The era of satellite limb observations started for more than 30 years ago with the launch of the Nimbus-6 satellite on 12 June 1975. The LRIR (Limb Radiance Inversion Radiometer, Gille et al. [1980]) instrument aboard this satellite had a water vapour channel in the infrared, but the calibration of this channel proved to be troublesome. First satellite-borne middle atmospheric water vapour results came from two instruments aboard the Nimbus-7 satellite launched three years later, namely the SAMS (Strato- spheric and Mesopheric Sounder, Drummond et al. [1980]) and the LIMS (Limb Infrared Monitor of the Stratosphere (LIMS), Gille et al. [1980]) instrument [Taylor et al., 1981; Fischer et al., 1981]. The SAMS instrument measured both emission and fluorescence in spectral bands around 2.7 µm and between 25 µm and 100 µm, while LIMS had a water vapour channel centred at 6.3 µm. The LIMS instrument was operational for seven months until May 1979, while the SAMS instrument went out of service first in June 1983. In the 1980s the development and construction of new instruments was pushed forward. In October 1984 ERBS (Earth Radiation Budget Satellite)

25 Chapter 2: Odin was launched, carrying aboard the SAGE II (Stratospheric Aerosol and Gas Experiment II, Mauldin et al. [1985]) instrument. This instrument used the solar occultation technique, which utilizes measurements of sunrise and sunset to measure attenuated solar radiation through the Earth’s limb. This technique is rather accurate, but with limited spatial and temporal coverage. SAGE II provided stratospheric water vapour, retrieved from a spectral channel centred at the 940 nm water vapour absorption band. Measurements lasted more than two decades until August 2005, providing the longest water vapour time series by a single satellite-borne instrument so far and benchmark results for stratospheric ozone and aerosols. A successor, SAGE III [Thomason and Taha, 2003], was launched in December 2001 but the measurement had to be terminated already in March 2006. The launch of UARS (Upper Atmosphere Research Satellite, Reber et al. [1993]) on 12 September 1991 heralded a mile stone of satellite-borne water vapour measurements in the middle atmosphere. Aboard the satellite were four instruments measuring water vapour at once, HALOE (Halo- gen Occultation Experiment, Russell et al. [1993]), MLS (Microwave Limb Sounder, Barath et al. [1993]), ISAMS (Improved Stratospheric and Meso- spheric Sounder, Taylor et al. [1993]) and CLAES (Cryogenic Limb Array Etalon Spectrometer, Roche et al. [1993]). Like SAGE II HALOE employed solar occultation, but used a spectral channel at 6.61 µm for the retrieval of water vapour. After more than a decade HALOE measurements ceased in the end of 2005, accomplishing major improvements in the understanding of middle atmospheric chemistry in general. The MLS instrument used for the first time thermal emissions in the microwave region. The water vapour product was retrieved from the 183 GHz emission line. Measurements finally ceased in July 1999. ISAMS and CLEAS were both only short-lived and measured emissions in the infrared. During the 1990s several instruments have been carried on space shuttle missions. The ATMOS (Atmospheric Trace Molecule Spectroscopy) instrument, a Fourier transform spectrometer, was carried by a space shuttle for the first time already in May 1985. The instrument flew on three more space shuttle missions in April 1992, April 1993 as well as in November 1994. These last three flights were part of NASA’s ATLAS (Atmospheric Laboratory for Applications and Science) series , i.e. ATLAS-1, -2 and -3 [Kaye and Miller, 1996]. The ATMOS instruments operated in solar occultation mode, acquiring high resolution infrared solar absorption spectra. The ATLAS series comprised also measurements of the MAS (Millimeter- wave Atmospheric Sounder, Croskey et al. [1992]) instrument, obtaining, like UARS/MLS, thermal emission spectra of water vapour at 183 GHz. The ATLAS-3 mission included in addition measurements of the CRISTA (Cryogenic Infrared Spectrometers and Telescopes for the Atmosphere, Of-

26 Chapter 2: Odin fermann et al. [1999]) and the MARSHI (Middle Atmosphere High Reso- lution Spectrograph Investigation, Conway et al. [1999]) instrument. Both instruments were mounted on the German SPAS (Shuttle Pallet Satellite) platform and deployed and retrieved by the space shuttle. CRISTA water vapour is primarily derived from emission spectra around 6.3 µm. MARSHI water vapour profiles are based on OH solar fluorescence emission spectra around 309 nm. One approach is to derive OH concentrations from these spectra, which are inverted to water vapour concentrations by means of a photochemical model [Summers et al., 1997, 2001]. Recently Stevens et al. [2008] presented a new approach, which allows the derivation of water vapour concentrations directly from the OH spectra, following the photolysis of water vapour by solar Lyman-α. In August 1997 CRISTA and MARSHI were put on a second mission, again on the SPAS platform [Grossmann et al., 2002; Conway et al., 1999]. The POAM II (Polar Ozone and Aerosol Measurements, Glaccum et al. [1996]) instrument and its follow-on POAM III [Lucke et al., 1999] were launched aboard the French SPOT 3 (Satellite Pour l’Observation de la Terre) in September 1993 and aboard SPOT 4 in March 1998 respectively. Both instruments employed the solar occultation technique and used like SAGE II the 940 nm water vapour absorption band for the water vapour retrieval. POAM II provided measurements until November 1996, POAM III until December 2005. Much more short-lived were the measurements of the Japanese solar occultation instruments ILAS (Improved Limb Atmospheric Sounder, Nakajima et al. [2002] and its successor ILAS II [Nakajima et al., 2006]. Their measurements lasted only from October 1996 to June 1997 and from April to October 2003 respectively. Water vapour profiles were based on spectral information in the infrared. Besides Odin currently five satellite missions provide middle atmospheric water vapour measurements. The TIMED (Thermosphere-Ionosphere-Meso- sphere Energetics and Dynamics) satellite is in orbit since December 2001. One of four instruments aboard is SABER (Sounding of the Atmosphere using Broadband Emission Radiometry, Russell et al. [1999]). Preliminary water vapour results derived from the 6.3 µm band have been presented recently. The MIPAS (Michelson Interferometer for Passive Atmospheric Sounding, Fischer et al. [2007]) instrument aboard the European Environ- mental Satellite (Envisat) operates since March 2002 and the retrieval of water vapour is based on 14 microwindows between 6 µm and 13 µm [Milz et al., 2005]. In August 2003 the Canadian ACE (Atmospheric Chemistry Experiment) satellite was launched. The FTS (Fourier Transform Spec- trometer, Bernath et al. [2005]) instrument operates in a solar occultation mode in the wavelength range from 2.2 µm to 13.3 µm. The water vapour retrieval employs 60 microwindows [Carleer et al., 2008]. The EOS MLS

27 Chapter 2: Odin

(Earth Observing System Microwave Limb Sounder, Waters et al. [2006]) instrument aboard the Aura satellite launched in July 2004 is the successor of UARS/MLS. Launched on 25 April 2007 AIM (Aeronomy of Ice Mission, Russell et al. [2008]) is the latest satellite to measure middle atmospheric water vapour. This satellite mission is dedicated to investigate noctilucent clouds. Water vapour measurements are performed by the SOFIE (Solar Occultation for Ice Experiment, Gordley et al. [2008]) instrument, using diﬀerential absorption measurements at 2.6 µm (strong absorption) and 2.45 µm (weak absorption). Over the last three decades all these satellite-borne measurements in combination with ground-based measurements have established a general picture of the water vapour distribution up to an altitude of 80 km. Still, the database in the mesosphere up to this altitude is rather limited. In the upper part of the middle atmosphere, i.e. in the altitude range between 80 and 100 km, only few measurements exist. Among other objectives Odin is to provide a valuable contribution and extension to the existing water vapour database in the middle atmosphere [Murtagh et al., 2002].

2.2 The Odin satellite

Odin is a Swedish-led small scale satellite mission in co-operation with Canada, France and Finland. From the beginning this mission involved an equal time-sharing between measurements of astronomy targets and observations at the atmospheric limb. After a decade of preparations the satellite was ﬁnally launched on 20 Febru- ary 2001 into a quasi-polar, sun synchronous and near-terminator orbit at an altitude of somewhat more than 600 km. This orbit provides a latitude coverage between 82◦S and 82◦N, when measuring along the orbit track. Since the end of 2004 Odin is pointing oﬀ-track during certain periods as permitted by sun angle constraints, e.g. in the southern hemisphere summer time, allowing the latitudinal coverage to be extended towards the poles. A single orbit takes about 97 minutes, thus Odin is orbiting the Earth about 15 times a day. Due to its sun-synchronous orbit the satellite can only perform measurements at two local times at a given latitude (one local time on the ascending node and a second local time on the descending node). In the beginning of the mission the satellite passed the equator at 18:00 local time on the ascending node and at 6:00 local time on the descending node. This local time coverage has changed somewhat during the Odin mission, since the orbit altitude gradually decreases due to atmospheric drag. At the time of writing, i.e. July 2008, the actual orbit altitude is 580 km and the equator passing time has shifted forwards by about 52 minutes.

28 Chapter 2: Odin

The satellite carries aboard two co-aligned instruments, OSIRIS (Optical Spectrograph and Infrared Imager System) and SMR (Sub-Millimetre Ra- diometer), whereby OSIRIS is just employed in the aeronomy measurement mode. The optical spectrograph measures scattered sunlight in the wavelength range between 280 nm and 800 nm with a 1 nm resolution. The infrared imager system operates at 1263 nm, 1273 nm, and 1530 nm respectively [Llewellyn et al., 2004]. The Sub-Millimetre Radiometer measures passively thermal emissions at the atmospheric limb using a 1.1 m tele- scope. It is actually composed of four tunable sub-millimetre radiometers covering several frequency ranges between 486 GHz and 581 GHz and one millimetre radiometer measuring around 119 GHz. Two autocorrelators and one acousto-optical spectrograph can be attached to any of the radiometers to detect the received signal [Frisk et al., 2003]. OSIRIS measurements provide mesospheric water vapour information using the same approaches as MARSHI described in the last section [Gattinger et al., 2006b, a; Stevens et al., 2008; Gattinger et al., 2008]. The Odin water vapour data used in this thesis are based on SMR measurements of the 557 GHz band. These measurements and the water vapour retrieval are described in the following sections.

2.3 SMR measurements of the 557 GHz band

The SMR instrument measures several atmospheric emission lines of water vapour and its isotopes. An complete overview can be found in paper (1) of this thesis. Stratospheric water vapour can be derived from measurements of the 488 GHz band. Measurements of the 557 GHz band provide water vapour informations in the mesosphere. Measurements of this band have initially been performed on 4 days per month. In April 2006 the measurement frequency was increased to 7 days. In May 2007 this frequency was changed once more to about 11 days per month, when measurements of the astronomy mode ceased. These measurements are part of stratosphere/mesosphere scans. In the beginning of the mission these scans covered the altitude range between 7 km and 100 km. Starting in 2003 these scans were extended up to an altitude of 110 km. Both in 2005 and 2006 this top altitude was actually frequently exceeded. Given a scanning velocity of 0.75 km · s−1 it takes more than two minutes to complete such a stratosphere/mesosphere scan. In total between 40 and 45 scans are performed per orbit. The eﬀective altitude resolution of the measurements is of the order of 3 km. This resolution is determined by a combination of the integration time for a single tangent height (1.85 s), the spectrometer read- out time and characteristics of the antenna. The horizontal resolution in

29 Chapter 2: Odin this altitude layer is characterised by the path length at the tangent height and is about 400 km. The horizontal sampling is given by the movement of the satellite, which is approximately 7 km · s−1 projected on the ground, and the time needed to perform the scan. In accordance the horizontal sampling is typically about one measurement per 1000 km. For the measurements of the 557 GHz band two of the four radiometers are used, either the 549A1 radiometer (covering 541 GHz to 558 GHz) or the 555B2 ra- 250 89.3 km diometer (covering 547 GHz to 200 564 GHz), using in total three [K] 150 B diﬀerent frequency conﬁgura-

T 100 50 tions. In either case autocorre- 0 lator number 1 is attached as 556.6 556.8 557 557.2 the signal receiving spectrometer. The studies in this thesis 250 69.8 km are based exclusively on data 200 taken with the frequency con-

[K] 150 B ﬁguration ”19”. This conﬁgu-

T 100 50 ration involves the 549A1 ra- 0 diometer and has been used 556.6 556.8 557 557.2 throughout the Odin mission. It is at the same time also the 250 49.9 km most optimised measurement 200 conﬁguration for the 557 GHz

[K] 150

B band, as described in paper

T 100 50 (2). In ﬁgure 2.1 typical 0 measurement spectra of the 556.6 556.8 557 557.2 557 GHz band are shown for diﬀerent tangent heights. The 250 water vapour emission line is 200 centred at 556.936 GHz, which

[K] 150

B is the strongest in the entire

T 100 GHz-region. In the limb direc- 50 29.6 km 0 tion this line is optically thick 556.6 556.8 557 557.2 in the centre up to an altitude Frequency [GHz] of about 70 km to 75 km. This Figure 2.1: Examples of the water vapour is why an absorption-like fea- emission line at 557 GHz at diﬀerent tan- ture can be observed at lower gent altitudes. tangent altitudes around the line centre. Still in the upper middle atmosphere this emission line exhibits a high intensity. In the second panel from below in ﬁgure 2.1 also an ozone emission line at 556.711 GHz

30 Chapter 2: Odin can be observed.

2.4 Proﬁle retrieval

The retrieval process describes the derivation of water vapour profiles (level-2 data) from the calibrated emission spectra (level-1 data, see figure 2.1). The water vapour retrieval from the 557 GHz band uses the Op- timal Estimation Method (OEM) [Rodgers, 2000; Baron et al., 2001]. This least-squares method gives the maximum likelihood solution combining statistical a priori knowledge on the variability of the searched parameters with the information provided by the measurement, using the associated errors as weights. The a priori knowledge of the searched parameters is used to stabilize the solution. The non-linearity of the 557 GHz band retrieval requires an iterative retrieval scheme. This scheme is based on a Newton and Levenberg-Marquardt method [Lautié et al., 2001]. Water vapour, temperature, hydrostatic pressure and ozone profiles are retrieved simultaneously from the 557 GHz band [Baron et al., 2001]. The altitude range in which water vapour can be retrieved depends on different factors. The measurement bandwidth of the SMR instrument of 800 MHz combined with the high optical thickness even in the line wings leads to a lower altitude limit of about 40 km for the water vapour retrieval. The upper altitude limit is determined by the signal-to-noise ratio and the water vapour concentration itself and lies usually above 100 km. The retrieved altitude resolution is in general 3 km. The statistical error for a single water vapour profile is of the order of 5 % to 15 % below 80 km. Above 90 km the statistical error can easily exceed 50 %. Hence in order to get sensible results averaging is necessary at these altitudes.

31 32 Chapter 3

Hygrosonde

Satellite measurements, like the one from Odin, provide a more or less global overview of different chemical species and other geophysical variables. In terms of spatial resolution these kind of measurements are rather limited. The horizontal resolution, expressed by the path length at the tangent height, is an integration usually over a distance of several hundred kilometres. In the vertical the resolution is in general not better than a kilometre, more often of the order of 2 km to 3 km. Hence, small scale and heterogeneous structures cannot be addressed by satellite measurements. For these kind of problems in-situ measurements are indispensable. Paper (5) reports about high resolution water vapour measurements from a balloon and sounding rocket in the vicinity of the polar vortex. The polar vortex edge prevents horizontal mixing between polar vortex and mid-latitude air. Given the characteristics of water vapour as a dynamical tracer and the subsidence of air inside the vortex, the water vapour profiles inside and outside the vortex will have a very different character, resulting in a distinct change of the water vapour distribution across the polar vortex edge. This edge also tends to feature filamentous structures and layering mainly due to irreversible planetary wave breaking, adding another aspect of small scale variability as observed across the polar vortex edge. The high-resolution water vapour measurements reported here have been performed during the Hygrosonde-2 campaign on 16 December 2001, a comprehensive set of experiments to measure the middle atmospheric water vapour concentration and the background state of the atmosphere in the vicinity of the polar vortex. In total, three such campaigns have been performed so far, with Hygrosonde-1 on 5 December 1994 and Hygrosonde-3 on 10 January 2005.

33 Chapter 3: Hygrosonde

3.1 Measurement principle

The principle behind these water vapour measurements is OH ﬂuorescence hygrometry [Kley and Stone, 1978; Khaplanov et al., 1996]. This technique employs active photodissociation of water vapour at wavelengths below 137 nm using a vacuum ultraviolet light source. This produces prefer- ably OH in its ground state (X2Π) according to reaction 1.4. A small, but signiﬁcant, fraction of this water vapour photodissociation results in electronically excited OH∗(A2Σ+), as described below.

∗ 2 + 2 H2O + h · ν → OH (A Σ ) + H( S)(λ < 137 nm) (3.1)

The excited hydroxyl radical subsequently returns to its ground state,

OH∗(A2Σ+) → OH(X2Π) + h · ν (3.2) emitting primarily in the wavelength range between 306 nm and 326 nm. Examples of such ﬂuorescence emission spectra are given in ﬁgure 3.1. The actual water vapour

−11 x 10 concentration CH2O can 8 6.4 hPa be derived from the in- 0.09 hPa tensity I of the emis- 6 sion according to equa-

4 tion 3.3.

2 I = Q·K ·CH2O (3.3)

Intensity [arbitrary unit] 0 In this equation is K a hygrometer specific con- −2 stant comprising vari- 310 315 320 325 330 Wavelength [nm] ous aspects, such as Figure 3.1: OH fluorescence emission spectra be- the photon flux of the tween 306 nm and 326 nm for different pressure light source, the geom- conditions. At 6.4 hPa the spectrum is quenched, etry and spectral re- at 0.09 hPa quenching does not play a role. The sponse of the detection spectra have been obtained by laboratory mea- system, as well as the surements at frost point conditions of -80◦C. quantum yield and water absorption cross section. This constant has to be determined by calibration efforts. The decay of the excited OH∗(A2Σ+) state is not only determined by fluorescence, but also by pressure-depending quenching. In the Earth’s atmosphere this process becomes important below about 60 km in this particular case. This is taken into account in equation 3.3 by Q, the quenching factor, which actually describes the relative contribution of emission and quenching as

34 Chapter 3: Hygrosonde competing processes responsible for the decay of OH∗(A2Σ+). For the Hy- grosonde instruments this factor has also been obtained by means of calibration measurements. More information can be found in [Gumbel et al., 1997].

3.2 Instrument aspects

The water vapour measurements require an uncontaminated measurement volume. For the in-situ hygrometry this is a major challenge, since out- gassing and desorption of water vapour from the payload itself can never be avoided completely. For the rocket-borne experiment the key to undisturbed measurements is to perform them outside the shock front, which keeps contaminations from the payload away from the measurement volume [Khaplanov et al., 1996]. Figure 3.2 describes this aero- dynamical concept visually. This concept works very well on the ascent, it is optimised for this leg of the rocket ﬂight. On the descent however the measurements are in general in- ﬂuenced by contaminations from out- gassing and desorption. Useful data will be only available at those altitudes where favourable payload atti- tudes allowed undisturbed measurements outside the shock front. The balloon-borne measurements might be contaminated on the ascent, when they are performed in the wake Figure 3.2: Design and aerodynam- of the balloon. During the relative ics of the rocket-borne hygrometer. fast descent of the balloon conditions are more optimised for the water vapour measurements. The instrument is looking downward to reduce contaminations and optimise the measurements further. The Hygrosonde instruments used a Lyman-α light source to initiate the photodissociation of the water vapour molecules. The Hygrosonde-1 rocket- borne instrument and the balloon-borne instruments used only one capillary

35 Chapter 3: Hygrosonde discharge inside the lamp body. For the Hygrosonde- 2 and -3 rocket-borne instruments six capillary discharges were comprised in the lamp body in order to improve the sensitivity of the hygrometer. The lamp is mod- ulated between on and off to discriminate the fluorescence signal from the background. Initially a symmetric modulation cycle was used. Since Hygrosonde-2 an asymmetric modulation cycle has been employed, with short intense pulses and longer periods for the background determination. This asymmetric cycle provides a better signal-to-noise ratio of the measurement, describing another effort to improve the sensitivity of the hygrometer. In total the Hygrosonde-2 rocket-borne instrument had an about 20 times higher sensitivity as compared to the Hygrosonde-1 instrument. The raw data have been read out every 2.5 ms during measurements from the rocket-borne Hygrosonde-2 instrument. Subsequently the data have been integrated over a period of 0.3 s in order to increase to the signal-to-noise ratio. The effective vertical resolution of these measurements is in order between 100 m to 300 m.

36 Chapter 4

Results of this thesis

• Paper (1)

The majority of the papers in this thesis are based on Odin/SMR water vapour data retrieved from measurements of the 557 GHz band. The first paper provides a general overview of the different emission lines of water vapour and its isotopes observed by Odin/SMR. The retrieval characteristics of these lines are described and examples of the global water vapour distribution retrieved from them are presented. Furthermore initial validation results are discussed. For the water vapour retrieved from the 557 GHz band this first validation showed a systematic dry bias in the lower mesosphere.

• Paper (2)

This paper followed up on the retrieval issues of the 557 GHz band noted in paper (1). The issues comprised also the temperature results retrieved from that band, which exhibited a cold bias in the entire mesosphere. In order to improve and resolve these problems a sensitivity study to identify the most critical parameters affecting the retrieval was performed. The sensitivity study was focused on polar summer mesosphere conditions, where the initial validation showed the high- est deviations to correlative measurements. The study revealed that uncertainties in the instrumental characterisation were significantly affecting the water vapour and temperature retrieval results. It was shown that also uncertainties in the broadening parameter of the emission line have a clear effect on the retrieved water vapour profile. In

37 Chapter 4: Results of this thesis

addition a the strong sensitivity of the retrieved water vapour concentration around 80 km to the tested parameters was notable. This result is important with respect to the water vapour peak formed due to the water redistribution by ice particles in the polar region during summer time.

• Paper (3)

The aim of this paper has been to characterise the mesospheric part of the semi-annual oscillation (SAO) in water vapour based on Odin/SMR measurements. This is an example of Odin’s water vapour observations revealing new inside into the dynamics of the middle atmosphere. These observations show a distinct phase shift of the SAO in the altitude range between 75 km and 80 km in the equatorial region. Higher up the water vapour maxima can be observed around the equinoxes, below this small layer the maxima occur around the solstices. The Odin/SMR observations show that the SAO is the dominant pattern of variability between 30◦S and 10◦N up to an altitude of 80 km. Above 90 km the SAO is dominating at all latitudes in the tropics and subtropics. Furthermore the measurements reveal that the annual component of the mesospheric water vapour variation shows distinct inter-hemispheric differences, being much stronger in the northern hemisphere subtropics as compared to its southern hemisphere counterpart. Also seasonal inter-hemispheric differences can be found in the Odin observations. In the altitude range between about 60 km and 90 km the northern hemisphere subtropics exhibit in general more water vapour in summer and less in winter as compared to subtropics in the southern hemisphere. One possible contribution to these observations might be differences in the meridional circulation from the summer pole towards the winter pole.

• Paper (4)

This paper presents Odin/SMR water vapour measurements in the upper mesosphere and lower thermosphere up to an altitude of 110 km. The emphasis was put on the polar area in winter time. Our observations show a distinct seasonal increase of the water vapour concentration during winter at a given altitude above 90 km. Above 95 km the observations exhibit the annual water vapour maximum during winter time. This behaviour is diﬀerent as observed in the subjacent part of the mesosphere, where the annual maximum occurs

38 Chapter 4: Results of this thesis

during summer time. Model simulations show similar results as the observations. Both, observations and model simulations show also clear inter-hemispheric differences in the water vapour distribution in the polar upper mesosphere and lower thermosphere during winter. The seasonal water vapour increase during winter is much more pronounced in the southern hemisphere, comprising higher concentrations and a larger latitude extent. We suggest that these differences are mainly due inter-hemispheric differences of the underlying dynamics.

• Paper (5)

This last paper focuses on water vapour measurements from balloon and rocket during the Hygrosonde-2 campaign on 16 December 2001. These measurements provided a detailed picture of the water vapour distribution in the vicinity of the polar vortex. Water vapour was used to determine the relative location of the measurements with respect to the polar vortex. The analysis showed that transitions between vortex and extra-vortex coincided with wind shears caused by gravity waves. A combination of the measurements from the Hygrosonde-1 and Hygrosonde-2 campaign provides a picture of typical water vapour concentrations that might be observed inside and outside the polar vortex. Even at an altitude of 70 km distinct diﬀerences between the vortex and extra-vortex water vapour concentrations can still be observed, clearly demonstrating that the vortex edge prevents horizontal mixing even at these high altitudes.

39 40 Chapter 5

Outlook

Even if Odin orbits the Earth’s atmosphere since more than seven and a half years the analysis of its middle atmospheric water vapour measurements is still rather in its beginnings. Hence many topics remain to be addressed. A primary objective is a continued retrieval optimisation and validation with correlative measurements. Continuous in-orbit test measurements to monitor and understand possible future changes in the instrument characterisation will be an important part of this objective. An important step will be the optimisation of the frequency configuration ”13”, another configuration used to measure the 557 GHz band. This configuration involves the 555B2 radiometer (see section 2.3) and has been initially used only during special measurement campaigns. Since April 2006 this configuration is used on a regular basis. With the new measurement schedule since May 2007 this configuration is used for the majority of the 557 GHz band measurements, i.e. usually 7 days out of 11 days per month. Of great interest is also a detailed comparison between the SMR and OSIRIS mesospheric water vapour results. On the scientific side the analysis of water vapour variability on different time scales will be of major interest. An interesting subject here will be signatures in the water vapour distribution originating from the QBO. As described in the introduction the QBO is a major pattern of variability and the effects of this oscillation can be found globally. For example the interannual variability of the mesospheric SAO described in paper (3) is very much determined by the QBO. Above 80 km the Odin/SMR observations show a steady increase over time in the water vapour concentration at all latitudes. The most likely expla- nation for this observed behaviour is the solar cycle influence on the water

41 Chapter 5: Outlook vapour [e.g. Chandra et al., 1997; Sonnemann et al., 2005]. In 2001 the solar cycle was on its maximum, resulting in a high flux of Lyman-α radiation and accordingly in a high water vapour photodissociation rate. Now, in 2008, the solar cycle is around its minimum, correspondingly leading to a lower photodissociation rate and higher water vapour concentrations. This analysis requires a detailed evaluation of the effects of the gradual change in the local time characteristics of the Odin orbit. Noctilucent clouds in the polar summer mesopause region have fascinated people ever since their first observation nearly 125 years ago. Due to their sensitivity to the ambient temperature and water vapour these clouds have been proposed and debated as a climate change indicator [e.g. Thomas et al., 1989; Thomas and Olivero, 2001; von Zahn, 2003; Thomas et al., 2003]. Colder mesopause temperatures due to increased levels of carbon dioxide and an increased water vapour concentration due to higher atmospheric concentrations of methane should favour the formation of these clouds. Ob- servations have indicated changes of NLC during the last decades in terms of brightness, occurrence number, season length and latitude extent [e.g. Klostermeyer, 2002; DeLand et al., 2006; Gadsden, 2002; Wickwar et al., 2002]. In addition inter-hemispheric differences have been noticed [e.g. Chu et al., 2001, 2006; DeLand et al., 2003; Bailey et al., 2005; Hervig and Siskind, 2006; Lübken and Berger, 2007]. While the temperature structure in the polar mesopause region is established, the water vapour concentration in this much more poorly known. Odin/SMR measurements of water vapour and temperature combined with Odin/OSIRIS measurements of cloud exis- tence and strength will certainly shed some more light onto the basic nature of these clouds, their redistribution of water vapour, their variability as well as their inter-hemispheric differences. Another intent is to deepen our understanding of the pole-to-pole circulation in the mesopause region. As described this circulation is governed by the angular momentum deposition from gravity waves, comprising upwelling in the mesosphere over the summer pole and downwelling over the winter pole. Of interest is not only the characterisation of the upward and downward transport in general or during anomalous events, but also the coupling of different parts of the middle atmosphere by this circulation. Recently evidence has come forward of an inter-hemispheric coupling between the polar winter stratosphere and the polar summer mesosphere, using the pole-to-pole circulation as the connecting link [Karlsson et al., 2007; Becker et al., 2004; Becker and Fritts, 2006; Karlsson et al., 2008]. Weaker (stronger) planetary activity in the polar winter stratosphere leads to a stronger and colder (weaker and warmer) polar vortex accompanied by a stronger (weaker) zonal flow. Changes in the zonal wind modulate the upward propagation of the gravity waves. As a consequence, stronger

42 Chapter 5: Outlook

(weaker) zonal flow will strengthen (weaken) the pole-to-pole circulation in the mesopause region, causing a stronger (weaker) upwelling in the polar summer mesosphere. Since the polar vortex and the zonal flow are gener- ally stronger in the southern hemisphere winter stratosphere, the northern polar summer mesopause should in principle exhibit colder temperatures and more water vapour. This causal connection appears to explain both inter-hemispheric differences and inter-annual variability of NLC. A stronger mesospheric pole-to-pole flow from the northern to the southern hemisphere might also account to a certain degree for the seasonal inter-hemispheric differences in water vapour discussed in paper (3) of this thesis. Tidal variations play an important role above 70 km and will always be embedded in the Odin/SMR data. The measurement database of the tidal variation in water vapour is very small. As described in section 2.2 Odin covers only two local times at a given latitude due its sun-synchronous orbit and is therefore not capable to provide detailed information on tidal variations. Also the capability of satellite-borne solar occultation measurements to extract tidal signatures is rather limited, since they are always performed at sunset and sunrise. Data from ground-based radiometers usually needs to be integrated over 24 hour period to provide reasonable results at 80 km. In the near future new ground-based efforts and measurements of the SABER instrument aboard the TIMED satellite may provide some new information. The TIMED satellite has a slowly precessing orbit, so that within 60 days all local times are covered at a given latitude. Also the combination Odin/SMR water vapour measurements with the results from Aura/MLS might be fruitful to gain some new knowledge of tidal signatures, at least in the tropical region, where the migrating diurnal tide is dominating. The Aura satellite has also a sun-synchronous orbit, passing the equator about 5 hours earlier than Odin at the moment. Also the implementation of results from the satellite-borne solar-occultation instruments, like ACE/FTS, UARS/HALOE and AIM/SOFIE, will be beneficial for this analysis. With the rather long data set available from Odin/SMR a final goal is to provide a unified middle atmosphere climatology from the different Odin/SMR water vapour products for reference and as input for model studies. This climatology will besides the data retrieved from the 488 GHz and 557 GHz band also include data derived from the 544 GHz band. This band is usually used for the retrieval of ozone and nitric acid (HNO3) and does actually not contain a clear emission line of water vapour. The water vapour information is contained in the spectral base line, which is part of the far wing of the strong 556.936 GHz water vapour emission line. This information allows to retrieve water vapour in the altitude range between about 15 and 35 km. There is also hope to extend the Odin/SMR water vapour data set a few kilometres beyond 110 km, based on the measurements where the top

43 Chapter 5: Outlook tangent height exceeded 110 km. In-situ high resolution water vapour in the middle atmosphere, like during the Hygrosonde campaigns, are very rare, but very much needed to address small scale structures and processes. Highly desired are for example high resolution water vapour measurements in the polar summer mesosphere, especially inside and in the vicinity of noctilucent clouds and PMSE to improve the understanding of their nature and microphysics. However the current design of the Hygrosonde instruments allow only measurements during nighttime conditions, avoiding solar background disturbing the measurements. There are several possibilities to modify the Hygrosonde instrument to achieve such desired measurements during daylight conditions. One possibility is to measure the OH resonance ﬂuorescence (see reactions 3.1 and 3.2) which is caused by the Sun, instead of using an active technique. This is the same principle as used by MARSHI and Odin/OSIRIS observations to derive water vapour (see section 2.1 and Stevens et al. [2008]). This approach requires higher spectrally resolved measurements as performed by the Hygrosonde instruments to resolve the spectral features needed for the water vapour retrieval. Another approach is based on active photodissociation of water vapour at wavelengths between 65 nm and 75 nm to produce excited atomic hydrogen:

∗ H2O + h · ν → OH + H (2p) (5.1) The excited hydrogen atom subsequently returns to its ground state H∗(2p) → H(1s) + h · ν (5.2) emitting Lyman-α radiation, which might be used to infer the water vapour concentrations. Such an instrument needs to look downward, since the sky above is illuminated by direct and scattered Lyman-α radiation disturbing the measurements. A completely diﬀerent possibility is to develop a rocket-borne radiometer measuring the thermal emission of water vapour, like the SMR instrument aboard Odin. For the forthcoming PHOCUS (Particles, Hydrogen and Oxy- gen Chemistry in the Upper Summer Mesosphere) rocket campaign, to be launched in 2010, such an instrument has been proposed. This experiment will actually consist of two radiometers. One radiometer will measure the 557 GHz band in a vertical viewing direction. Diﬀerentiation of the emissions at two altitudes can be used to derive the water vapour concentration in the altitude region in-between. A second radiometer will measure the 183 GHz water vapour emission line in sideway direction, similar to satellite limb measurements. Early ideas for such rocket-borne radiometer measurements have been described by Croskey et al. [1994].

44 Acknowledgements

Being used to five year plans, this part of my life comes to an end actually ahead of schedule, adding some more grey hairs, wrinkles and also some extra kilograms due to the excessive consumption of chocolate and cookies to sustain the overall pressure. Many people have accompanied me during this time and contributed in very different ways and I would like to acknowledge them here. First and foremost I would like to thank the wonderful people in the AP group. I will keep you in my heart for the rest of my life. I especially enjoyed the nice times we had together during all the travels, campaigns and dinners. Thanks to...... Jörg Gumbel, who made this journey possible in the very first place. A good friend and an inspiring and thoughtful supervisor. I think he was suffering sometimes quite a lot with four Ph.D. students at the same time. ...Linda Megner, who hardly ever managed to be in the office before me but then always showed up with a big smile greeting me with ”Guten Morgen, Herr Lossow!” I certainly will miss that! ...Bodil Karlsson, ”die gelbe Schlange”, who never forgot any of the stupid German phrases I taught her but always the practical ones and who had to hold out with me during our trips to Canada. ...Jonas Hedin, who shared his office with me during the last months and who was exposed to too many German speaking people. At the end he was away quite often, but I hope there is no connection to that. ...Jacek Stegman, for organising several workshop trips for me including all the travel grants! ...Farahnaz Khosrawi, Miss ”particularly suitable” who disrupted the current at the Esrange hotel, while trying to get out a slice of bread from a toaster with a fork. Otherwise she always did a very good job in commenting my manuscripts. ...Heiner Körnich for being my favourite opposing player during all floor hockey games. Thanks for all the help and discussions, sometimes you could be really nice to me! ;-) ...Misha Khaplanov, who patiently tried to explain me all technical details

45 Acknowledgements of the Hygrosonde measurements. I’m sure I will find some more questions! ...Georg Witt for sharing his incredible knowledge, ideas, and all stories from the 1960s and 1970s. ...Tomas Waldemarsson, who unfortunately left a group much too early, but in that way opened up the possibility to move into Jonas’ office. Many thanks to Admir Targino, my first room mate at MISU, with who I had a very nice time together. A huge thank to all Ph.D. students and to all the other people at MISU for creating such a nice and relaxed atmosphere in this institute. Thank you all! A special thank to our administrative per- sonal for handling all kind of stuff, especially complicated travel expenses. I’m grateful to Birgitta Björling for quickly taking care of all my publica- tion requests, which were mainly written before I was born or published in journals I never had heard of before. Outside of MISU I would like to thank my fellow co-authors at Chalmers, i.e. Jo Urban, Donal Murtagh and Patrick Eriksson. Especially the position of Jo in these acknowledgements does not display by any means his importance for this thesis. He has without any doubt deserved to be named at the very first place!!! Thanks for sharing your knowledge and all your help!!! I’m very grateful for it! Also a thank to Nicolas Lautié who introduced me into the Odin water vapour measurements and its retrieval. He shares the same faith into Matlab as myself. Only Pete knows why! Thanks to Ted Shepherd for inviting me to the University of Toronto and letting me play with CMAM data. I’m sorry that there was not enough time to finish our project and make it part of this thesis. I really appreciate the help I got from Andreas Jonsson and Charles McLandress while working on this project! At this point I also would like to take the opportunity to thank both the people at the Department for Radio and Space Science (Chalmers University of Technology in Göteborg) and at the Department of Physics (University of Toronto) for the nice and friendly atmosphere experienced there during my stays. My appreciation also to Karl Heinrich Fricke for the instructive time at Esrange. In good old tradition I would like to thank also the artists who were accom- panying me musically through this time. Thanks to DJ André Tanneberger, alias ATB, and The Prodigy (Thanks to my sister for the singles collection, I guess you do not want to have a copy?). Thanks also to Die Toten Hosen, a childhood favourite, which rather unexpected re-entered my life. Thanks for that to my relative Ingo, playing them in his car on a early Sunday morning in the spring of 2006 heading towards Berlin-Schönefeld Airport. Also thanks to my teammates in various Korpen football teams, for all the fun we had and making me forget the work for a while. We might be even more successful if we would play it the German way.

46 Acknowledgements

Finally I would like to thank my family for their support through all this time. A bigger thank than you ever can imagine to my wife Ulrike for your incredible support and patience. For a person who is rather impatient this a major accomplishment!!! I think that says all about your commitment! ILD!

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